Giovanni Aloisi

Keywords: sulfur cycle, ocean biogeochemistry, salt giant, gypsum, Mediterranean Sea, Messinian Salinity Crisis, multiple sulfur isotopes, iron sulfide, sulfate Chapter 1-General Introduction IPGP Communication Service

5.8.1 First scenario: Pore water sulfur isotopic distillation as a driver for anomalously high isotopic composition of iron sulfides? ………………………………..………………………..230 5.8.2 Scenario 2: concurrent influence of variable fractionation factors and sulfide pooling…………………………………………………………………………………………………………………….238 vi 5.8.3 Scenario 3: influence of variable sulfur fractionation factors in a virtually infinite pool of sulfate …………..…………………………………………………………………………………………..243 5.8.4 Scenario 4: Influence of re-oxidative processes ………………………………….………246 5.9 Conclusion …………………………………………………………………………………………………………..254 Chapter 6. Conclusion and perspectives………………………………………………………………….……259 Bibliography………………………………………………………………………………………………………..………273 Appendix……………………………………………………………………………………………………………………..307 Appendix A. DSDP and IODP sulfur and oxygen isotopic data measured in this study ……….307

Résumé

La Crise de Salinité Messinienne (MSC) est un évènement géologique et paléocéanographique autant exceptionnel que controversé ayant eu lieu à la fin du Miocène . La MSC a profondément affecté le bassin Méditerranéen, notamment parce qu'il a donné lieu à la déposition d'un immense volume d'évaporites : le Géant Salifère de Méditerranée (MSG). Ce bouleversement sans précédent de la Méditerranée est le résultat de la diminution graduelle des échanges hydrologiques entre la Méditerranée et l'océan Atlantique, qui a éventuellement permis l'isolation complète du bassin Méditerranéen.

De nombreuses preuves suggèrent qu'avant même le début de la précipitation des évaporites, la biogéochimie de la Méditerranée a été profondément impactée par la restriction des échanges hydrologique entre la mer et l'océan. Cependant, le fonctionnement biogéochimique des bassins méditerranéens profonds et marginaux durant la MSC est peu connu, principalement parce que les couches de sels qui constituent le MSG sont dépourvues des marqueurs couramment utilisés pour les reconstructions paléo-océanographiques.

Dans ce manuscrit, j'aborde cette question ouverte en étudiant le cycle biogéochimique du soufre dans les couches constituées de sulfate de calcium déposées lors de la MSC. Le cycle du soufre est particulièrement intéressant dans ce contexte car i) il est un des éléments principaux formant le gypse et l'anhydrite, minéraux qui constituent une partie importante du géant salifère et ii) le cycle du soufre est particulièrement sensible aux perturbations biogéochimiques induites lors d'une restriction hydrologique. Les modifications du cycle du soufre sont par conséquent étroitement liées aux cycles des éléments majeurs comme le carbone. Afin d'identifier et de quantifier les processus redox qui ont contrôlé le cycle du xiv soufre dans les bassins marginaux et profonds de la Méditerranée pendant la MSC, j'ai utilisé les isotopes multiples du soufre (δ 34 S et ∆ 33 S) comme outil principal.

Cette approche m'a permis d'établir un lien entre la présence de fossiles de bactéries sulfoxidante dans le gypse de bassins marginaux de la Méditerranée et un cycle cryptique du soufre impliquant la réduction du sulfate en sulfure puis la re-oxidation quasi quantitative de ce sulfure en sulfate. Cette dynamique biogéochimique particulière a impliqué la reminéralisation d'importants flux de carbone organique benthiques (3.8mm.cm -2 .ans -1 ) par la sulfato-réduction. Ces flux sont comparables à ceux mesurés dans les zones les plus productives de l'océan moderne où des communautés bactériennes similaires à celles fossilisées dans le gypse du MSG participent activement à un cycle cryptique du soufre.

Dans la deuxième partie de mon doctorat, j'ai utilisé une approche similaire afin de pouvoir décrypter le cycle du soufre dans les bassins profonds et ouverts de la Méditerranée à la fin de la MSC. D'une part, cette approche m'a permis de comprendre que le cycle cryptique du soufre n'était pas restreint aux bassin marginaux et peu profonds de la Méditerranée. Il était aussi largement impliqué dans le cycle du soufre des bassins profonds à la fin de la MSC. D'autre part, les signatures isotopiques des sulfures de fer reflètent souvent l'influence de paramètres sédimentaires locaux sur le microenvironnement dans lequel ils se forment. Dans cette thèse, je propose que l'épaisseur de la colonne d'eau sous laquelle les sulfures de fer se forment joue un rôle majeur sur leur composition isotopique. J'ai donc utilisé les signatures isotopiques des sulfures de fer contenus dans les dépôts évaporitiques des bassins profonds de la Méditerranée pour déterminer la profondeur maximale des bassins dans lesquels les sulfures de fers se sont formés à la fin de la MSC. Cette estimation a ensuite été utilisée par une de mes collègue doctorante comme paramètre lors des reconstructions paléobathymétrique des bassins Méditerranéens, permettant ainsi de contraindre la baisse xv minimale de niveau marin ayant pu affecter la Méditerranée durant la MSC. Les résultats de cette étude suggèrent une baisse de niveau marin d'au moins 1100m à l'Ouest du seuil de Sicile et de 2500m à l'Est du seuil de Sicile. Ces résultats valident un modèle de déconnection hydrologique des deux principaux bassins de la Méditerranée à la fin de la MSC, comme l'ont précédemment suggéré des études géophysiques et géologiques indépendantes.

Mots clés: cycle du soufre, biogéochimie de l'océan, géant salifère, gypse, Mer Méditerranée, Crise de salinité Messinienne, isotopes multiples du soufre, sulfures de fer, sulfate xvii

Résumé étendu

Les environnements évaporitiques, tels que ceux qui ont caractérisé les bassins méditerranéens lors de la crise de salinité messinienne (MSC), sont souvent imaginés comme des déserts, hostiles pour la plupart des formes de vie. La Mer « Morte », souvent considérée comme plus proche analogue moderne du géant salifère de Méditerrannée en est un bon exemple. Néanmoins, alors que les évaporites qui caractérisent la Méditerranée pendant la MSC sont exempts de macro-fossiles, ce travail de thèse a permis de montrer qu'un cycle biogéochimique du soufre intense, impliquant une importante activité microbienne, avait lieu lors du dépôt du gypse qui constitue le géant salifère de Méditerranée. J'ai utilisé les isotopes multiples du soufre dans les minéraux contenant des sulfates et des sulfures comme outil principal pour révéler la nature du cycle biogéochimique du soufre. Pour cela, je me suis concentrée sur les dépôts constitués d'alternances cycliques entre marne et sulfate de calcium (gypse/anhydrite), qui sont caractéristiques à la fois de la phase initiale (Primary Lower Gypsum, PLG) et de la phase terminale (Upper Gypsum, UG) de la MSC.

Ces alternations lithologiques sont interprétées comme reflétant des changements de l'hydrologie des bassins méditerranéens (marginaux et profonds), eux même contrôlés par les oscillations climatiques induites par les cycles astronomiques de précession (Andreetto et al., 2022;Lugli et al., 2010 ;[START_REF] Manzi | A new facies model for the Upper Gypsum of Sicily (Italy): chronological and paleoenvironmental constraits for the Messinian salinity crisis in the Mediterranean[END_REF]. Pour la première fois nous avons montré que ces changements hydrologiques (oscillant entre plus humides et plus arides) contrôlaient également la signature isotopique des sulfures de fer préservés dans les dépôts évaporitiques et les marnes. Les sulfures de fer contenus dans les marnes ont en effet des compositions appauvries en 34 S tandis que ceux extrait du gypse sont plus variables et en moyenne enrichis xviii en 34 S. Nous avons interprété ces résultats en utilisant une approche qui inclut la modélisation numérique de l'évolution de la composition isotopique du soufre en fonction des différents flux et processus qui peuvent l'affecter.

Grâce à cette démarche, nous avons démontré que les différentes gammes de δ 34 S mesurées correspondent à des dynamiques redox du cycle du soufre distinctes pendant les périodes de précipitation du gypse et de déposition des marnes. Les sulfures de fer qui présentent les valeurs les plus élevées en δ 34 S et les plus faibles en ∆ 33 S correspondent à ceux contenus dans les gypses et anhydrites. Ces compositions isotopiques ont été interprétées comme étant le résultat d'un cycle du soufre lors duquel (i) les fractionnements isotopiques exprimés pendant la sulfato-réduction étaient relativement faibles, et (ii) le sulfure d'hydrogène produit par sulfato-réduction était quasiment quantitativement ré-oxydé, résultant en un cycle cryptique du soufre. En comparaison, la composition isotopique des sulfures de fer dans les marnes caractérisés par des valeurs de δ 34 S plus faibles et maximales de ∆ 33 S. Cela suggère que le cycle du soufre pendant le dépôt des marnes était caractérisé par une ré-oxydation du sulfure d'hydrogène moins complète et par de plus grands fractionnements isotopiques exprimés lors la sulfato-réduction.

Dans les dépôts de PLG formés dans les bassins marginaux de la Méditerranée au début de la MSC, nous avons interprété ces différentes tendances comme reflétant l'alternance entre un cycle cryptique du soufre pendant la formation du gypse et de manière générale, des flux moins importants de carbone et de soufre impliqués dans les processus de sulfato-réduction and de re-oxydation du sulfure d'hydrogène lors de la déposition des marnes. D'une part, ces conclusions supportent les observations d'études indépendantes qui ont suggéré précédemment que les marnes alternant avec le gypse des dépôts PLG, se sont déposées dans xix des bassins stratifiés ayant des colonnes d'eau épisodiquement euxiniques (e.g., Sinninghe Damsté et al., 1995 ;Isaji et al., 2019). D'autre part, le cycle cryptique du soufre pendant la période de formation du gypse qui constitue le PLG était contrôlé par des consortiums de bactéries utilisant le soufre, incluant des sulfo-oxydantes dont les restes ont été perminéralisés dans le gypse. Ces consortiums sont fonctionnellement similaires à ceux retrouvés au fonds des marges continentales modernes très productives où un cycle cryptique du soufre est aussi observé (Bruchert et al., 2006 ;[START_REF] Treude | Rapid sulfur cycling in sediments from the Peruvian oxygen minimum zone featuring simultaneous sulfate reduction and sulfide oxidation[END_REF]Callbeck et al., 2021). Le fait que plus de 80% du soufre ultimement incorporé dans la formation du gypse ait été d'abord impliqué dans un cycle cryptique du soufre supporte le rôle important des bactéries sulfo-oxydante dans la formation du gypse dans les bassins marginaux de la Méditerranée pendant le stade 1 de la MSC. Cet aspect est plus particulièrement développé dans l'étude de Aloisi et al. (2022) à laquelle j'ai contribué (Appendice C).

Nos résultats dans les dépôts des phases 1 et 3 de la MSC (PLG et UG) démontrent que la variabilité des compositions isotopiques mesurées dans les sulfures de fer préservés était largement contrôlée par leurs environnements de dépôt.

De nombreux auteurs ont précédemment suggéré que les variations eustatiques sont un des facteurs important contrôlant ces paramètres environnementaux (e.g., [START_REF] Fike | Rethinking the ancient sulfur cycle[END_REF]Houghton et al., 2022;Pasquier et al., 2021a;Pasquier et al., 2021b;Pasquier et al., 2017). De fait, les variations du niveau marin ont le potentiel d'exercer une influence majeure sur les compositions isotopiques des sulfures de fer puisque la profondeur de la colonne d'eau, et la distance vis-à-vis des sources continentales d'un bassin donné, contrôlent largement les paramètres environnementaux tels que les taux de sédimentation, les flux benthiques de matière organique, les flux de ré-oxydation des sulfures et le mélange vertical des sédiments xx (Aller, 2010;Aller, 2014;Dunne et al.,2007;[START_REF] Jørgensen | Sulfur biogeochemical cycle of marine sediments[END_REF]Middelburg et al., 1997), qui a leur tour impactent la dynamique des micro-environnements dans lesquels les sulfures de fer se forment.

Dans cette thèse, j'ai utilisé ces observations pour estimer le niveau d'eau des bassins méditerranéens pendant le stade terminal de la MSC, qui reste une période particulièrement controversée pour les scenarii paléo-environementaux de la MSC. J'ai couplé mon approche isotopique à des observations pétrographiques. Celles-ci m'ont permis de déduire que les sulfures de fer contenus dans les dépôts de la phase 3 ont été formés pendant la diagénèse précoce. Par la suite, une compilation des valeurs isotopiques mesurées dans des sédiments modernes et anciens a permis de confirmer qu'il existe une corrélation entre la profondeur du bassin dans laquelle se forment les sulfures de fer pendant la diagénèse précoce et leurs compositions isotopiques. En général, des valeurs isotopiques plus faibles et moins variables sont retrouvées dans les bassins excédant 1.5 km de profondeur. Considérant que les compositions des sulfures de fer dans les dépôts de la phase 3 ont des δ 34 S particulièrement variables, et en moyenne enrichis en isotopes lourds, nous avons donc suggéré que ces sulfures se sont déposés dans des bassins dont les profondeurs n'excédaient pas 1.5 km. Nous proposons que ces sulfures de fer ont été formés à des profondeurs moins importantes que celles auxquelles ils ont été échantillonnés (1965 à 4450 mbsl).

En collaboration avec Hanneke Heida, également doctorante de l'ETN SaltGiant, j'ai utilisé cette profondeur maximale de dépôt des couches évaporitiques afin de contraindre l'amplitude de la baisse du niveau de la Méditerranée à la fin de la MSC. Les résultats de cette approche suggèrent une baisse du niveau de leau de 1.1 km de la Méditerranée occidentale et de 2.5 km de la Méditerranée orientale. Cette découverte suggère ainsi une déconnection xxi de ces deux bassins à la fin de la MSC, comme le suggèrent d'autres études indépendantes (e.g., Heida et al., 2022 ;[START_REF] Madof | Discovery of vast fluvial deposits provides evidence for drawdown during the late Miocene Messinian salinity crisis[END_REF][START_REF] Camerlenghi | Seismic markers of the Messinian salinity crisis in the deep Ionian Basin[END_REF][START_REF] Lofi | Erosional processes and paleo-environmental changes in the Western Gulf of Lions (SW France) during the Messinian Salinity Crisis[END_REF].

Finalement, mes travaux de thèse ont permis de souligner que la variabilité dans la composition isotopique des sulfures de fer dans le milieu sédimentaire peut être, dans certains cas, largement attribuée aux variations de deux facteurs : (i) les flux realtifs de sulfures produits par sulfato-réduction par rapport à ceux, et subséquemment re-oxydés en composés intermédiaires et/ou en sulfate (ii) le facteur de fractionnement lors de la sulfatoréduction. Alors que ces deux facteurs sont souvent considérés comme fixes et/ou négligeables lors de reconstructions du cycle rédox du soufre, la prise en compte de leurs variations au cours du temps, a des implications importantes pour une compréhension plus précise du cycle du soufre et plus largement pour mieux comprendre comment la biogéochimie de l'océan répond aux perturbations environnementales, autant dans les environnements modernes que passés.

Chapter 1.

General introduction 1.1 Sulfur cycle in the ocean

The tightly intertwined biogeochemical cycling of sulfur and carbon in Earth surface reservoirs (sediments, ocean and atmosphere) is both a consequence and a cause of climate and Earth oxidation state changes (Crockford et al., 2019;Ebelmen., 1845;Lyons et al., 2014).

Organic carbon formed by photosynthetic CO2 fixation in the upper ocean sinks to the ocean floor and is oxidized back into carbon dioxide via biogeochemical reactions involving a sequence of electron acceptors. This results in the sub-division of the zone of organic matter remineralization into sub-zones dominated by a distinct oxidative process (aerobic respiration, denitrification, sulfate reduction…). This subdivision occurs because oxidants of organic matter are utilized successively according to their energy yield, starting from O2 (which provides the greatest energy yield), followed by NO3 -, Fe-and Mn-oxides, sulfate and, finally, the fermentation of organic matter, which is the incomplete oxidation of organic matter in the absence of external electron acceptors (Figure 1. 1) (Froelich et al., 1979). Sulfate (SO4 2-) -sulfur in its most oxidized form (oxidation state +VI) -is the second most abundant anion in seawater, about 150 times more concentrated than O2, and thus represents the largest oxidant pool in the ocean (Hayes and Waldbauer, 2006). Hence, sulfate plays a fundamental role in oceanic carbon cycle. Jørgensen and Kasten (2006), modified from Froelich et al. (1979). Note that the methane (CH4) and sulfate (SO4 2-) concentration curves cross each other at the lower limit of the sulfate reduction zone (Jørgensen and Kasten, 2006).

Organoclastic microbial sulfate reduction (MSR) (Equation 1) is responsible for between 12 and 65% of the remineralization of the organic carbon flux to the seafloor [START_REF] Bowles | Global rates of marine sulfate reduction and implications for sub-sea-floor metabolic activities[END_REF][START_REF] Bradley | Widespread energy limitation to life in global subseafloor sediments[END_REF]Jørgensen and Kasten, 2006), whereas SO4 2-consumption via the anaerobic oxidation of methane (AOM) (Equation 2) acts as a barrier for CH4 emissions from the subsurface (Figure 1. 1) (Boetius et al., 2000;Egger et al., 2018;Knittel and Boetius, 2009;Treude et al., 2005).

Organoclastic sulfate reduction: SO4 2-+ 2CH2O → H2S + 2HCO3 - (1) Methane based sulfate reduction: CH4 + SO4 2-→ HCO3 -+ HS -+H2O

(2)

Most of the hydrogen sulfide produced by MSR and AOM diffuses vertically and is rapidly reoxidized back into sulfate, further consuming oxidants, and is so prevented from diffusing into the water column [START_REF] Liu | Early diagenesis of iron and sulfur in Bornholm Basin sediments: The role of near-surface pyrite formation[END_REF]Fossing and Jørgensen, 1990;[START_REF] Jørgensen | Sulfur biogeochemical cycle of marine sediments[END_REF]. Furthermore, when sulfide-reoxidation is driven by microbial activity, it has the potential to significantly participate to dark carbon fixation (Dyksma et al., 2016).

Of the H2S produced by MSR, a fraction is scavenged by reactive iron to form solid iron sulfides (the most reduced form or sulfur with oxidation state -II) that are ultimately buried in marine sediments, the prevalent mineral phase involved in this process being pyrite (FeS2) (Equations 3,4,5) [START_REF] Rickard | Sulfidic sediments and sedimentary rocks[END_REF].

Fe 2+ + H2S → FeS +2H + [START_REF] Rickard | Chemistry of Iron Sulfides[END_REF]

And FeS + H2S → FeS2 +H2 (H2S pathway) [START_REF] Rickard | Kinetics of pyrite formation by the H2S oxidation of iron (II) monosulfide in aqueous solutions between 25 and 125°C: The rate equation[END_REF]

Or FeS + Sx 2-→FeS2 + Sx-1 2-(Polysulfide pathway) [START_REF] Rickard | Chemistry of Iron Sulfides[END_REF] (5)

The above-mentioned biogeochemical processes make redox sulfur cycle a major determinant of the earth surface oxidation state (Turchyn and Schrag, 2004): the burial of reduced sulfur species acts as a net source of oxygen to the earth surface (Berner, 1987); while, once emerged and exposed to weathering, the oxidation of pyrite is a sink of atmospheric oxygen (Equation 6):

Pyrite oxidation: FeS2 + 15/4O2 +7/2H2O → Fe(OH)3 + 2H2SO4 (6)

Weathering of pyrite also has an effect on the carbon cycle because it produces sulfuric acid (Equation 6) that further participates in continental weathering processes and might: (i) decrease the alkalinity produced during silicate weathering, decreasing carbonate deposition, and (ii) act as a net source of CO2 to the atmosphere when weathering carbonates because it does not involve the utilization of atmospheric CO2 in the weathering process (Equation 7 and 8) (Bufe et al., 2021;Burke et al., 2018;Calmels et al., 2007;Kölling et al., 2019;Torres et al., 2014).

Carbonate weathering using sulfuric acid: 2CaCO3 + H2SO4 → 2Ca 2+ + 2HCO3 -+ SO4 2- (7) Carbonate weathering using CO2: CaCO3 + CO2 + H2O → Ca 2+ + 2HCO 3- (8) Finally, the biogeochemical state of the ocean is also affected by the sequestration of calcium and sulfate ions from the ocean through the deposition of calcium sulfate minerals (mainly gypsum and anhydrite) resulting from the evaporitic concentration of seawater (Halevy et al., 2012;Wortmann and Chernyavsky, 2007;[START_REF] Wortmann | Rapid variability of seawater chemistry over the past 130 million years[END_REF]) (Equation 9):

Gypsum precipitation: SO4 2-+ Ca 2+ + 2H2O → Ca.SO4.2H2O (9)

First, large scale marine deposition of calcium sulfate causes the removal of the oxidant SO4 2-from the surface system [START_REF] Prince | Evaporite deposition in the mid-Neoproterozoic as a driver for changes in seawater chemistry and the biogeochemical cycle of sulfur[END_REF], making the ocean more sensitive to external perturbations in O2 and CO2 because it modifies the sedimentary fluxes of pyrite and organic carbon (Halevy et al., 2012;[START_REF] Wortmann | Rapid variability of seawater chemistry over the past 130 million years[END_REF]. Second, the extraction of oceanic dissolved Ca 2+ results in a decrease in the global calcium carbonate deposition flux. The consequential rise in alkalinity could lead to the pumping of atmospheric CO2 by the ocean, a weaker greenhouse effect and global cooling (Shields and Mills, 2020). Conversely, evaporite weathering supplies calcium to the ocean, enhances carbonate precipitation which is a source of CO2, strengthens greenhouse forcing and leads to a warmer climate (Shields and Mills, 2020) (Figure 1. 2). Carbonate weathering (green) releases calcium and alkalinity from the rock reservoir and CO2 from the atmosphere to the ocean. Carbonate deposition (black) reverses this process and decreases alkalinity and calcium concentration and releases CO2 to the atmosphere. Silicate weathering (purple) transfers calcium from the rock reservoir to the ocean and two moles of carbon from the atmosphere to the ocean. Evaporite weathering (orange) transfers only calcium to the ocean. From Shields and Mills (2020).

The various processes introduced above -redox sulfur cycling, iron sulfide and evaporite burial and weathering -occur on timescales spanning from days (sedimentary redox cycling) (Fike et al., 2009;Gomes et al., 2022;Jørgensen, 1977;Kristensen et al., 2000;Mengdi et al., 2022;[START_REF] Seitaj | Cable bacteria generate a firewall against euxinia in seasonally hypoxic basins[END_REF] to hundreds of thousands of years and beyond (continental weathering and S-mineral burial) (Canfield, 2004;Hay et al., 2006;Leavitt et al., 2013;Lécuyer, 2016;Masterson et al., 2016;Paytan et al., 1998;Turchyn and Schrag, 2004;[START_REF] Wortmann | Rapid variability of seawater chemistry over the past 130 million years[END_REF]Wu et al., 2010). In addition to the time scale considered, their impact on ocean chemistry and carbon cycling depends on the degree of connectivity of the environment they act in with the global oceanic sulfate reservoir (Andreetto et al., 2022a;[START_REF] Fike | Rethinking the ancient sulfur cycle[END_REF]) -I will expand on time scales and connectivity below. Fossing and Jørgensen (1990); [START_REF] Jørgensen | A Thiosulfate Shunt in the Sulfur Cycle of Marine Sediments[END_REF]; [START_REF] Thamdrup | Thiosulfate and Sulfite Distributions in Porewater of Marine-Sediments Related to Manganese, Iron, and Sulfur Geochemistry[END_REF].These redox processes occur mostly in small restricted marine reservoirs, generally in the porewater environment, but can also occur within the water column in specific cases. In any case, these restricted reservoirs have limited exchanges with the global ocean (blue arrow). The fluxes of sedimentary sulfide and sulfate minerals being buried in the rock record (brown arrows) are the net result of the redox processes. While the orders of magnitude for the fluxes shown are well known, their precise estimation depends on the method used to calculate, as well as on the time period for which the flux is integrated and if the oceanic sulfur cycle is considered at steady state. a corresponds to the estimated flux of sulfate burial, considering that no major evaporitic basins are actually precipitating sulfate bearing salts and that barite burial (∼0.02 Tmol.years-1) is negligible compared to the fluxes considered here (Carter et al., 2020;Dickens et al., 2003) b Estimated fluxes considering sulfur cycle at steady state and integrating the flux of SO4 buried into sulfate bearing minerals over the course of the last 20Myrs (Tostevin et al., 2014) C Estimated flux after (Berner, 1982) d,e,f,g Range of existing estimates of annual fluxes of sulfate being reduced by MSR via various methods [START_REF] Bowles | Global rates of marine sulfate reduction and implications for sub-sea-floor metabolic activities[END_REF][START_REF] Bradley | Widespread energy limitation to life in global subseafloor sediments[END_REF]Canfield et al., 2005;Thullner et al., 2009). h,i,j Reoxidative fluxes of hydrogen sulfide to sulfate calculated using estimates from d,e,f and following Fossing and Jørgensen (1990); [START_REF] Jørgensen | A Thiosulfate Shunt in the Sulfur Cycle of Marine Sediments[END_REF]; [START_REF] Thamdrup | Thiosulfate and Sulfite Distributions in Porewater of Marine-Sediments Related to Manganese, Iron, and Sulfur Geochemistry[END_REF], that consider that 90% of the sulfate being processes by MSR is ultimately reoxidized back to sulfate k Reoxidative fluxes of hydrogen sulfide to sulfate considering that numbers given in references b,c,d,e,f represent valid estimations and steady state i.e., the flux of sulfur that is not buried as sedimentary sulfide is necessarily reoxidized. In that case, the 90% of sulfide reoxidation estimation, widely used in the literature (see h,i, jfor references) is not valid anymore for the upper range of sulfate reduced by MSR estimation and up to 99% of the sulfide produced by MSR should have been reoxidized to satisfy mass balance. bearing minerals over the course of the last 20Myrs (Tostevin et al., 2014) l Estimated fluxes of sulfur from the weathering of sulfur bearing minerals (Burke et al., 2018) m,n Net fluxes of sulfur resulting from both volcanic inputs and hydrothermal activity outputs from Alt (1995);and Kagoshima et al. (2015) O corresponds to the amount of sulfate dissolved in the ocean (Petsch and Berner, 1998) In contrast to the small reservoirs where redox cycling of S occur, sulfate in the global ocean is a large reservoir (4.10 19 mol of dissolved SO4 2-;Petsch and Berner (1998)) and has a very long residence time (10 7 years) (Holser and Kaplan, 1966;Lécuyer, 2016)(Figure 1. 4). Further, compared to this long residence time, the mixing time of the ocean is relatively rapid (∼2000 years, obtained dividing ocean volume of 1.37 x 10 18 m 3 by the Global Ocean Conveyor Belt water flux of approximately 20 x 10 6 m 3 /s, (Broecker, 1992)), leading to a spatially homogeneous oceanic sulfate concentration.

The net outgoing fluxes of sulfur from the restricted reservoir where redox cycling occurs (sulfate and sulfide bearing minerals burial) (Figure 1. 3) affect the long-term S cycle (Figure 1.4). In turn, long term changes in the fluxes of oxidized and reduced sulfur species to and from the large oceanic reservoir (volcanic emissions, hydrothermal reactions, continental weathering, sedimentary S burial)

control the oceanic sulfate concentration on time scales of several million years, which feeds back on ocean biogeochemistry and S redox cycling (Berner, 1989;Ebelmen., 1845;Garrels, 1984)(Figure 1. 4).

A crucial issue is that of the relative importance of the sedimentary, biogeochemical iron sulfide flux versus the sedimentary, evaporitic calcium sulfate flux. While iron sulfides form continuously in the MSR and AOM zones of marine sediments [START_REF] Rickard | Sulfidic sediments and sedimentary rocks[END_REF], marine calcium sulfate evaporite deposition depends on the occurrence of hydrologically restricted marine environments where seawater is concentrated by evaporation [START_REF] Warren | Evaporites, Sediments, Resources and Hydrocarbons[END_REF].

In addition, sulfur is one of the main components of evaporitic minerals whereas iron sulfide constitutes at maximum a few percent of the sedimentary rocks in which it is preserved. Finally, the rate of precipitation/weathering of calcium sulfate generally exceeds that of other sedimentary rocks that contains pyrite (Nichols et al., 2009).

In summary, the modification of the oceanic sulfate concentration, and the related changes in ocean biogeochemistry, depend on a complex interplay of tectonic, eustatic and climatic factors acting on geological timescales. These factors control both short-term, episodic changes in the precipitation and weathering rates of calcium sulfate minerals, as well as long-term changes in the large and continuous burial and oxidation fluxes of sedimentary sulfide (Halevy et al., 2012;[START_REF] Wortmann | Rapid variability of seawater chemistry over the past 130 million years[END_REF]. For example, episodic drops in the global ocean sulfate concentration have been inferred to be due to rapid and massive extraction of sulfate ions from the oceanic reservoir caused by evaporites deposition (Hay et al., 2006). In conjunction with other factors such as important should be accumulating in sediments porewater. Such large accumulation of H2S and intermediate S species or sedimentary sulfide are not observed and are way larger than the most generous estimates for sedimentary sulfide burial rates (0.6 to 1.2 Tmol.years -1 ) (Berner, 1982;Tostevin et al., 2014;Wortmann and Chernyavsky, 2007). On the other hand, if the upper range estimates for yearly microbial sulfate reduction hold true (as suggested in the Jørgensen (2021) review), estimates of the global rate of sedimentary pyrite burial imply a nearly quantitative re-oxidation of sulfide globally (Figure 1. 3). This question regarding the modern sulfur cycle remains unanswered and requires to be further investigated in order to understand the biogeochemistry of sulfur in aquatic environments.

Reconstructions of the past oceanic sulfur cycle are even more problematic since there is no direct access to biogeochemical rates, and direct estimates of ancient sulfate concentrations from fluid inclusions are scarce and rely on the occurrence of evaporitic deposits that might not even always reflect the composition of seawater since they form in basins that are restricted with respect to the global ocean (Brennan et al., 2013;Horita et al., 2002;[START_REF] Lowenstein | Secular variation in seawater chemistry and the origin of calcium chloride basinal brines[END_REF]Spear et al., 2014).

To tackle the S cycle in the geological past, the more common approach developed since the late 60's is the measurement of the sulfur isotopic composition (δ 34 S) of minerals bearing sulfur in its oxidized (sedimentary sulfate) or reduced (sedimentary sulfide) forms [START_REF] Canfield | Biogeochemistry of sulfur isotopes[END_REF]Canfield, 2004;Canfield et al., 2000;[START_REF] Claypool | The age curves of sulfur and oxygen isotopes in marine sulfate and their mutual interpretation[END_REF]Holser and Kaplan, 1966;Paytan et al., 1998;Thode and Monster, 1965;Wu et al., 2010).

This approach is possible because the concentration of sulfate and its sulfur isotopic composition are controlled by the same set of S fluxes. MSR produces sulfide depleted in 34 S compared to the source oceanic sulfate (Fry et al., 1986;Thode and Monster, 1965;[START_REF] Thode | Sulphur Isotope Geochemistry[END_REF]. If the MSR and sulfide burial fluxes increase, both the reduced mineral sulfur and the residual sulfate isotopic composition evolve towards higher values, following a Rayleigh type isotopic distillation process. For this reason, by using Rayleigh calculations, and measuring sulfur isotopic composition of sulfate-and sulfur-bearing minerals, it is theoretically possible to trace back the changes in oceanic sulfate concentration and the relative magnitudes of S cycle fluxes. Specifically, the enrichment of sedimentary S-bearing minerals in 34 S are often interpreted to reflect a decrease in sulfate concentration and the long-term increase of the relative amount of sulfate consumed by MSR and buried as reduced sulfur (e.g. Canfield, 2004;Halevy et al., 2012;[START_REF] Wortmann | Rapid variability of seawater chemistry over the past 130 million years[END_REF]Yao et al., 2018).

Nevertheless, as seen in section 1.1 (Figure 1.3), sulfur and sulfide bearing minerals form in reservoirs that have restricted exchanges with the global ocean (see section 1.1 and Figure 1. 3). Recent studies have highlighted the role of the local environment in controlling the isotopic composition of sulfur bearing minerals [START_REF] Fike | Rethinking the ancient sulfur cycle[END_REF]Houghton et al., 2022;Jones and Fike, 2013;Marin-Carbonne et al., 2022;Pasquier et al., 2021a;Pasquier et al., 2021b;Pasquier et al., 2017). This is because local sedimentary or environmental settings influence biological settings and often determine the dynamics of redox sulfur cycling in these reservoirs. These observations challenge classical interpretation of sulfur isotope composition. Yet, while former approaches were focused on the quantification of the net fluxes of solid sulfur out of these restricted reservoirs over long timescales, the conundrum above provides an opportunity to estimate the fluxes of sulfur in the restricted reservoirs where redox sulfur cycling occurs.

The isotopic composition of S burial fluxes is controlled by the rate at which sulfur is reduced by MSR, re-oxidized or sequestered as solid sulfide or sulfate. However, in most of the studies attributing changes in δ 34 S sulfur isotopic composition to fluctuations in local environmental conditions, the changes in sulfate concentration produced by MSR in porewaters, leading to the accumulation of 34 Senriched iron sulfides (i.e., sulfide pooling), is the main proposed control of the isotopic composition of sedimentary sulfides (Houghton et al., 2022;Liu et al., 2021;Pasquier et al., 2021b;Pasquier et al., 2017). This is because (i) sulfate is easily depleted in porewater environments (ii) sulfate minerals don't precipitate spontaneously in normal marine environments. Thus, the interpretations are often performed only considering the variations in the isotopic composition of sedimentary sulfides. These studies, however, neglect supplementary sources of S isotope variability: (i) the effect of changes in cell specific rates of sulfate reduction (csSR) and other environmental factors (e.g., temperature, sediment mineralogy or type of strains) on the S isotope fractionation between source SO4 2-and product H2S (ii) the effect of sulfide re-oxidative processes that can overprint the isotopic signature initially produced by MSR [START_REF] Findlay | Sulfide oxidation affects the preservation of sulfur isotope signals[END_REF][START_REF] Liu | Early diagenesis of iron and sulfur in Bornholm Basin sediments: The role of near-surface pyrite formation[END_REF]Leavitt et al., 2013;Mengdi et al., 2022). The means by which sulfur isotopic fractionation during MSR, and sulfide reoxidation, affect sulfur isotopic composition will be further developed in the result chapters of this manuscript (Chapter 3, 4 and 5).

There are several ways to assess the influence of these poorly investigated controls on the sulfur fluxes and isotopic composition of sedimentary S. On the one hand, the measurement of the minor sulfur isotopes 33 S/ 32 S (in addition to 34 S/ 32 S), as well as the 18 O/ 16 O ratio of sulfate-bearing minerals have the potential to help deconvolve the relative influence of sulfide re-oxidative processes occurring concomitantly to MSR (e.g., Böttcher et al., 2001;Johnston et al., 2007;Johnston et al., 2005;Torfstein and Turchyn, 2017;[START_REF] Turchyn | Oxygen isotopic composition of sulfate in deep sea pore fluid: evidence for rapid sulfur cycling[END_REF][START_REF] Zerkle | Fractionation of multiple sulfur isotopes during phototrophic oxidation of sulfide and elemental sulfur by a green sulfur bacterium[END_REF]. On the other hand, environments with high sulfate concentrations -where both sulfate and sulfide minerals form concomitantly -provide rare opportunities to deconvolve the relative effect of sulfide pooling from other controls on sulfur isotopic composition, because in sulfate-rich environments the Rayleigh distillation that should affect both sulfate and sulfide isotopes in that case should be measurable (Jones and Fike, 2013).

Despite the fact that sulfur is a major player in oceanic biogeochemistry and is one of the building blocks of marine evaporites, its (bio)geochemical cycling during the formation of large bodies of marine sulfate bearing minerals such as gypsum or anhydrite has been overlooked. Studying both

Chapter 1-General Introduction _________________________________________________________________________________________________ 14 the sulfate and sulfide multiple sulfur isotopic composition of minerals formed coevally during evaporitic events provides an opportunity to quantitatively understand the processes that control sulfur isotope compositions in the restricted reservoirs where sulfur redox processes take place. The Earth surface has sporadically experienced periods of massive deposition of evaporitic minerals in marine basins. These so-called marine salt giants are characterized by their impressive volumes (covering between 10^5 to 10^6 km 2 and with thicknesses often over 1km) and were generally deposited in tectonically-originated, hydrologically restricted oceanic marginal basins during periods of accretion or breakout of continents (Hay et al., 2006;Hsu¨, 1972;Warren, 2010) (Figure 1. 5).

Marine salt giants

In this regard, marine salt giants have no present day or recent equivalents because Quaternary evaporites were generally deposited in endorheic hypersaline lakes formed within intracontinental depressions [START_REF] Warren | Evaporites, Sediments, Resources and Hydrocarbons[END_REF] and have volumes that are one to four orders of magnitude smaller than those of marine salt giants of the geological record (Warren, 2010). Compared to the mass of salts contained in salt giants of the geological record, the initial amounts of ions extracted from the ocean through salt giant formation was most likely considerably higher, given that the evaporitic salts that constitute salt giants are easily eroded and dissolved [START_REF] Wortmann | Rapid variability of seawater chemistry over the past 130 million years[END_REF]. For example, the amount of calcium sulfate precipitated during the Neoproterozoic to early Cambrian have been estimated to represent only a quarter to what have been initially

extracted from the global ocean [START_REF] Wortmann | Rapid variability of seawater chemistry over the past 130 million years[END_REF]) (Figure 1. 6). Evaporitic salts that constitute marine salt giants are deposited via the evaporative concentration of sea water dissolved ions in a basin that is geographically marginal -and hydrologically restricted -from the global ocean.

In addition, the hydrological freshwater budget of the marginal basin has to be necessarily negative, i.e. the amount of water extracted from the basin by evaporation is larger than that added by continental runoff and precipitation [START_REF] Warren | Evaporites, Sediments, Resources and Hydrocarbons[END_REF]. When the evaporating brine is purely marine and has experienced limited mixing with non-marine waters, the succession of evaporitic minerals that precipitate chemically reflects the relative abundance of dissolved ions in sea water (Hardie, 1991). The chemical precipitation of a mineral from a saline solution occurs when its ionic product overcomes its solubility, that is to say, when the solution is sursaturated with respect to the mineral specie considered (Bąbel and Schreiber, 2014). The degree of saturation with respect to a given evaporitic mineral depends on the concentration that the solution acquired through evaporation (Harvie and Weare, 1980). Evaporitic minerals crystallize in a sequence [START_REF] Usiglio | Analyse de l'eau de la Mediterrane ´e sur les co ˆtes de France[END_REF], for which the order depends on the solubility of the minerals and thus, on the concentration of the solution (Harvie and Weare, 1980). The sequence of evaporitic minerals that form during the evaporation of sea water is shown in Figure 1. 7.

First, most of the bicarbonate ions are precipitated together with calcium to form carbonate. Then, gypsum (Ca.SO4.2H2O) precipitates and extracts calcium and sulfate from the brine at salinities of ∼110g.kg -1 and halite (NaCl) starts precipitating at salinities exceeding ∼320g.kg -1 (Harvie and Weare, 1980) (Figure 1. 7). Finally, further freshwater extraction from the brine leads to the formation of magnesium and potassium bitter salts (Harvie and Weare, 1980;[START_REF] Usiglio | Analyse de l'eau de la Mediterrane ´e sur les co ˆtes de France[END_REF].

Because salt giants are formed by evaporative concentration of sea water where sodium and chloride ions represent ∼85% by weight and calcium and sulfate ∼8% of all the dissolved species ( Figure 1. 8), halite (NaCl) is the dominant mineral constituting marine salt giants, followed by calcium sulfate (gypsum, CaSO4.2H2O and anhydrite, CaSO4) minerals (Hay et al., 2006;[START_REF] Warren | Evaporites, Sediments, Resources and Hydrocarbons[END_REF]. Lofi (2018) In this work, I investigated the Mediterranean salt giant (MSG), one of the youngest and bestpreserved marine salt giant on earth [START_REF] Hsü | Origin of Saline Giants: a critical review after the discovery of the Mediterranean evaporite[END_REF]. The MSG contains in total ∼1 million km 3 of evaporites, (mainly halite and gypsum) for which ∼95% are nowadays buried in deep basins of the Mediterranean below thick Plio-quaternary sediments (Haq et al., 2020;Hübscher et al., 2007;[START_REF] Ryan | Quantitative evaluation of the depth of the western Mediterranean before, during and after the late Miocene salinity crisis[END_REF] (Figure 1. 9). The remaining 5%, mainly in the form of gypsum, were mostly formed in marginal perched basins and are outcropping onshore due to post-depositional tectonically-induced uplift [START_REF] Decima | Late Miocene evaporites of the Central Sicilian Basin[END_REF]Hübscher et al., 2007) (Figure 1. 9). Overall, the formation of this salt giant is thought to have contributed to the extraction of 6 to 10% of the total dissolved ions from the global ocean (Garcia-Castellanos and Villaseñor, 2011;Ryan, 1973). The formation of the MSG is the result of a major tectonic-paleoenvironmental-oceanographic event known as the Messinian salinity Crisis (5.97 to 5.33Ma). The MSC was triggered by the partial to complete hydrological restriction of the connection between the Mediterranean Sea and the Atlantic ocean (Ryan, 2009). This catastrophic event occurred at the end of the Messinian stage (7.25 to 5.33Ma), and marks the transition from the Miocene (23. 03 to 5.33 Ma) to the Pliocene (5.33 to 2.58Ma) epochs (Krijgsman et al., 1999;Manzi et al., 2013;Roveri et al., 2014a).

Despite the identification of gypsum layers formed at the end of the Miocene in Sicily already in the late 19 th century (Mayer-Eymar, 1867), further assignation of similar onshore deposits to the same evaporitic event were not performed before the late 1950's to late 1960's in Sicily [START_REF] Ogniben | Petrografia della serie solfifera siciliana e considerazioni geologiche[END_REF]Selli, 1964), Algeria (Perrodon, 1957), Spain [START_REF] Rios | Saline deposits of Spain[END_REF] and Cyprus (Gass and Cockbain, 1961).

Simultaneously, the beginning of sea floor exploration through seismic studies allowed to detect the existence of extensive erosion features on Mediterranean submarine margins and of a major discordance at the base of the Pliocene in the deep basins of the Mediterranean (Bourcart, 1958(Bourcart, , 1959(Bourcart, , 1960)). Together with the discovery of offshore Mediterranean structures resembling to salt domes found in the Gulf of Mexico (Hersey, 1965;Menard. et al., 1965), and of a similar fossils assemblage across multiple sections outcropping in marginal areas of the Mediterranean [START_REF] Decima | Ostracodi del Gen. Cyprideis JONES del Neogene e del Quaternario italiani[END_REF]Ruggieri, 1967), this further suggested that important hydrologic and environmental changes, possibly leading to extensive erosion, evaporites precipitation and development of hyposaline environments occurred during the late Miocene and affected an important area of the Mediterranean (Ruggieri, 1967). This developing MSC theory was confirmed by the first Mediterranean scientific drilling expedition on the Glomar Challenger (DSDP project), that drilled across the above-mentioned unconformity detected on the seismic reflection data, confirming the occurrence of evaporitic deposits underlying Plio-Quaternary sediments over large areas of the Mediterranean (Ryan, 1973;Ryan and Cita, 1978).

Following this major discovery, an intense scientific effort developed to decipher the complex Chapter 1-General Introduction _________________________________________________________________________________________________ 20 intertwining of geological, hydrological, and geochemical processes that have led to the formation of this Mediterranean salt giant. This effort is still ongoing, with the number of publications containing the terms "Messinian salinity crisis" in their titles, abstract or key words increasing since then (Web of science, Scopus) (Figure 1. 10). The variety of study areas, analytical tools and scientific approaches led to the birth of several scenarios and controversies within the scientific community studying the Mediterranean salt giant.

The three original MSC scenarios are summarized in Roveri et al. (2014a) and are differentiated based on the physical depth (bathymetry) of Mediterranean basins and on the thickness of the water column that existed during the formation of the salt giant (Figure 1. 11 andFigure 1. 12).

The first model, « shallow basin, shallow water » suggests that the Mediterranean basins were shallow before and during the MSC, allowing the formation of evaporites under a reduced water column [START_REF] Nesteroff | Un modéle pour les évaporites messinienne en Méditerranée, bassins peu profonds avec dépot d'évaporites lagunaires[END_REF]. This scenario requires that after the MSC vertical tectonic movements during the Pliocene would have led to the subsidence of Mediterranean basins to their modern depth of 3000-4000 meters [START_REF] Nesteroff | Un modéle pour les évaporites messinienne en Méditerranée, bassins peu profonds avec dépot d'évaporites lagunaires[END_REF]. However, the occurrence of deep pelagic sequences below, within and above the evaporitic layers rapidly discarded this hypothesis (Hsü et al., 1973;[START_REF] Ryan | Quantitative evaluation of the depth of the western Mediterranean before, during and after the late Miocene salinity crisis[END_REF]. Hsü et al. (1973) and the bottom one represents "deep water, deep basins" [START_REF] Andreetto | Alternate interpretation of the Messinian salinity crisis, controversy resolved?[END_REF]. Modified from Krijgsman et al. (2018) Figure 1. 12 Two (simplified) conceptual scenarii derived from the two step model of [START_REF] Andreetto | Alternate interpretation of the Messinian salinity crisis, controversy resolved?[END_REF] and accounting for thick MSC halite deposition in the deep basin. Stage 1 (5.96-5.60 Ma, Krijgsman et al., 1999): primary gypsum accumulated in silled marginal basins. Stage 2: Scenario 1: moderate base level drop in the Mediterranean and evaporite emplacement by density currents in a stratified deep water environment [START_REF] Roveri | The Mediterranean Messinian salinity crisis: an Apennines foredeep perspective[END_REF]2009;2014b); Scenario 2: drying out of the Mediterranean basin highly concentrated in salts (brine pool) by evaporative shrinking [START_REF] Ryan | Modeling the magnitude and timing of evaporative drawdown during the Messinian salinity crisis[END_REF]Lofi et al., 2011b). Evaporite emplacement starts in deep water environment and ends in shallow water environment; During the MSC, continental margins and slopes underwent intense subaerial/subaqueous erosion. In red, the Messinian Erosion Surface (MES) where margins and slopes were subaerially exposed. The advocates of scenario 2 proposed that evaporites formation also occurred under a reduced water column, but in basins that were deep, due to a severe water level fall of the Mediterranean Sea: it is the "deep dessicated basin" scenario (Hsü et al., 1973) (Figure 1. 11, top panel; Figure 1. 12, scenario 2). In the third scenario, "deep water, deep basins" the Mediterranean basins are suggested to be deep as well, with no or very limited shallowing of the water column and evaporites either precipitated in deep waters or redeposited in deeper settings [START_REF] Andreetto | Alternate interpretation of the Messinian salinity crisis, controversy resolved?[END_REF] ( Figure 1. 11,bottom panel;Figure 1. 12,scenario 1). The evidence in support of either the desiccated or the full-basin Mediterranean models have now been studied for about 50 years at both onshore and offshore locations. However, these observations are still very difficult to reconcile with one another (Andreetto et al., 2021a;Raad et al., 2021). the Paratethys (from ∼11 to 6.9Ma). Modified from Palcu et al. (2021) Until the end of the Tortonian, (∼7.2Ma), the Atlantic Ocean and the Mediterranean Sea water exchanges were not occurring only, like today, through the narrow (∼13km width) strait of Gibraltar,

Chapter 1-General Introduction _________________________________________________________________________________________________ 24 but through the Betic (South Spain) and Rifian (North Morocco) corridors (Flecker et al., 2015;Roveri and Manzi, 2006) (Figure 1. 14). These seaways gradually disappeared throughout the Messinian Stage as a result of the convergence between the African and Eurasian plates coupled with subduction dynamics in the Alborán region [START_REF] Capella | Mediterranean isolation preconditioning the Earth System for late Miocene climate cooling[END_REF]Duggen et al., 2004) (Figure 1. 14).

Figure 1. 14 Paleogeographic evolution of the Mediterranean-Atlantic gateway from the Late Tortonian and early Messinian-before the MSC onset-to the Zanclean at the end of the MSC. From Capella et al. (2017), modified after [START_REF] Cornée | The Messinian erosional surface and early Pliocene reflooding in the Alboran Sea: new insights from the Boudinar basin[END_REF]Martín et al., 2009;[START_REF] Mather | A Field Guide to the Neogene Sedimentary Basins of the Almería Province[END_REF]Martín et al., 2014) Starting from 7.17Ma, the impact of the hydrological restriction caused by the closure of the Betic and Rifian corridors is observed in the sedimentary record of the Eastern Mediterranean through a shift from oxic to suboxic biotic assemblages (Kouwenhoven et al., 2003;[START_REF] Kouwenhoven | Paleoenvironmental evolution of the eastern Mediterranean during the Messinian: constraints from integrated microfossil data of the Pissouri Basin (Cyprus)[END_REF]Seidenkrantz et al., 2000) and the more frequent occurrence of marine sapropels (organic rich sediments formed under deoxygenated bottom water) (Santarelli et al., 1998;Seidenkrantz et al., 2000)(Figure 1. 15). Bulian et al. (2022), adapted from (Alhammoud et al., 2010;Pinardi and Masetti, 2000;[START_REF] Topper | The precessional phase lag of Messinian gypsum deposition in Mediterranean marginal basins[END_REF] In addition, a global aridification of the Mediterranean region seems to have occurred concomitantly to the hydrological restriction (Butiseaca et al., 2022). From this period onward, due to hydrological restriction, the Mediterranean basins became more sensitive to changes in the hydrological regime and oceanic circulation induced by orbital forcing (Bulian et al., 2022;Bulian et al., 2021) (Figure 1.

15).

From 6.8Ma, stratification and increased bottom water salinities are suggested to have occurred in both deep eastern and western basins of the Mediterranean (Drinia et al., 2007;Kouwenhoven et al., 1999;Pellegrino et al., 2020;Sierro et al., 2003;Zachariasse et al., 2021). In more peripheral-and thus-shallower areas of the Mediterranean, the drastic restriction of the Atlantic gateways led to the alteration of faunal morphologies [START_REF] Mancini | Calcareous nannofossil size and abundance response to the Messinian salinity crisis onset and paleoenvironmental dynamics[END_REF], to the burial and preservation of high amounts of organic matter (Dela Pierre et al., 2012;Kontakiotis et al., 2021;Tzevahirtzian et al., 2022) and to a drop in diversity and almost complete disappearance of calcareous plankton [START_REF] Krijgsman | The onset of the Messinian salinity crisis in the Eastern Mediterranean (Pissouri Basin, Cyprus)[END_REF]. Precession maxima was often characterized by the deposition of carbonate-rich layers associated to microbial activity, implying the occurrence of sulfate reduction and/or sulfide oxidation (Borrelli et al., 2021;Dela Pierre et al., 2012;Natalicchio et al., 2019;Tzevahirtzian et al., 2022) (Panels A and B, Figure 1. 16). Conversely, in intermediate to shallow basins of the Mediterranean, the lipid biomarker content and isotopic compositions suggest that the water column was intermittently stratified during precession minima, with occurrence of euxinia and phytoplankton blooms (Natalicchio et al., 2019;Sabino et al., 2022;[START_REF] Sabino | The response of water column and sedimentary environments to the advent of the Messinian salinity crisis: insights from an onshore deep-water section (Govone, NW Italy)[END_REF]Vasiliev et al., 2019) (Panels A and B, Figure 1. 16).

These conditions suggest the possible existence of a biogeochemically active S cycle in Mediterranean marginal basins prior to the onset of the MSC. Although our period of interest lies within the MSC time interval, these observations set the stage, biogeochemically, for the investigations presented in this thesis.

Timing and chronology of the Messinian salinity crisis events

Increasing hydrological restriction eventually led, at 5.97Ma, to the onset of the MSC. Starting from the MSC onset, the timing of successive MSC events is controversial due to the lack of chronostratigraphic markers and stratigraphic gaps caused by the erosion of MSC sedimentary units (Roveri et al., 2014b). Although there is no irrefutable evidence supporting a given MSC model proposed in the literature (see section 1.4 above), a "consensus" model was developed in 2008, (and later refined by Roveri et al. (2014a)) that attempted to consider the largest possible amount of existing evidences, and highlighting the questions that are still to be answered regarding the MSC [START_REF] Ciesm | The Messinian Salinity Crisis from mega deposits to microbiology[END_REF]. This "consensus" model allowed to develop a common vocabulary to name the different MSC stages and deposits. In the following manuscript, I will use this nomenclature that I detail below.

Following this model, the MSC is divided in 3 main stages during which the "trilogy" of evaporitic lithologies deposited in the western Mediterranean basin (gypsum-halite-gypsum) while mainly halite formed in the eastern Mediterranean (Lofi et al., 2011;Roveri et al., 2014a).

During the first stage of the MSC (5.971-5.60 Ma), the Primary Lower Gypsum unit (PLG), made of bottom grown gypsum beds alternated with marls was deposited in marginal, intermediate to shallow, water basins of the Mediterranean, from southern Spain in the west to Cyprus in the east (Lugli et al., 2010) (more details on PLG gypsum will be developed in Chapters 2 and 3 of this manuscript). In more proximal domains of marginal basins, evaporitic-microbialitic limestone (Terminal Carbonate Complex TCC or Calcare di Base CdB) where possibly deposited concomitantly to PLG; however the timing of formation of these carbonate beds (concomitantly to PLG or in more advanced stages of the MSC) is controversial (Bourillot et al., 2010;Cornée et al., 2004;Rouchy and Caruso, 2006;Roveri et al., 2008;Tzevahirtzian et al., 2022) (Figure 1. 17). In deeper settings of Mediterranean marginal basins, organic rich sediments and carbonates deposited, suggesting the inhibition of gypsum precipitation in the deep part of the basin, possibly caused by microbial sulfate consumption under a stratified water column (de Lange and Krijgsman, 2010;Dela Pierre et al., 2012;Natalicchio et al., 2019) (Figure 1. 17). Whereas the PLG mentioned above deposited in shallow, marginal basins of the Mediterranean at depths of 800m maximum (Ochoa et al., 2015;Raad et al., 2021), the nature of MSC deposits formed in the main deep basins of the Mediterranean during MSC phase 1 are poorly constrained because only seismic data (and no drilling), is available.

Mediterranean deep-sea MSC deposits possibly formed during MSC phase 1 are collectedly termed the Lower Unit (LU) (Lofi et al., 2011) (Figure 1. 17).

Figure 1. 17 Conceptual stratigraphic model of Messinian deposits from Roveri et al. (2014b). LU, MU and UU corresponds to the seismic units Lower Unit, Mobile Unit and Upper Unit respectively. MES, BES, BS, TS and TES correspond to erosional surfaces, M/P to the Miocene/Pliocene boundary.

Following PLG and LU deposition, the second stage of the MSC (5.60-5.55 Ma) is characterized by the intensification of Mediterranean hydrological restriction. This led to the precipitation of up to 3km of halite (on average 1km in the deep western basin and 2km in the deep eastern basin) (Mobile Unit, MU in (Bertoni and Cartwright, 2005;[START_REF] Gvirtzman | Intra-Messinian truncation surface in the Levant Basin explained by subaqueous dissolution[END_REF]Haq et al., 2020;Lofi et al., 2011) Figure 1. 17). This period is also characterized by intense erosion on the margins of the Mediterranean with incision of canyons up to 1500m deep in the Rhone and Nile valleys (Clauzon, 1974), and by the accumulation of the erosion product, including the products of eroded PLG called Resedimented Lower Gypsum (RLG) in the deep domains of the Mediterranean [START_REF] Lofi | Erosional processes and paleo-environmental changes in the Western Gulf of Lions (SW France) during the Messinian Salinity Crisis[END_REF] (Figure 1. 17).

During the third and terminal phase of the MSC (5.55-5.33 Ma), the Mediterranean freshened and the Upper Gypsum (UG) unit (or Upper Unit UU, in deep Mediterranean basins) was deposited. This unit is principally made of gypsum and marls interbeds -occasionally bearing brackish water ostracods -that overlie halite deposits in intermediate to abyssal depths of the Mediterranean basin (Andreetto et al., 2021a;Andreetto et al., 2022a;Andreetto et al., 2022b;Andreetto et al., 2021b;[START_REF] Manzi | A new facies model for the Upper Gypsum of Sicily (Italy): chronological and paleoenvironmental constraits for the Messinian salinity crisis in the Mediterranean[END_REF]Manzi et al., 2016;[START_REF] Rouchy | Paleoenvironmental changes at the Messinian-Pliocene boundary in the eastern Mediterranean (southern Cyprus basins): significance of the Messinian Lago-Mare[END_REF]Roveri et al., 2014b) (Figure 1. 17). The water column thickness and the overall connectivity of the various sub-Mediterranean basins during stage 3 of the MSC are amongst the most debated issues within the MSC scientific community. This will be further developed in chapters 2, 4 and 5 of this manuscript and is extensively explained in the review paper in appendix D. Finally, "normal" marine conditions were restored at ∼5.33Ma (Roveri et al., 2014a;[START_REF] Van Couvering | The base of the Zanclean Stage and of the Pliocene Series[END_REF]. The timing and vertical amplitude of the Mediterranean reflooding is also greatly debated (e.g., [START_REF] Andreetto | Impact of pulsed Atlantic water inflow into the Alboran Basin at the time of the Zanclean flooding[END_REF][START_REF] Garcia-Castellanos | Catastrophic flood of the Mediterranean after the Messinian salinity crisis[END_REF]Marzocchi et al., 2016), since this also depends on the state of the Mediterranean water level at the end of the phase 3 and thus on the preferred MSC scenario.

Rationale and scientific objective of the thesis

In addition to the prominent role of salt giants formation on oceanic sulfate, carbon cycling and earth surface oxidation on a long timescales (Myr) (e.g., Halevy et al., 2012;Shields and Mills, 2020;Wortmann and Chernyavsky, 2007;[START_REF] Wortmann | Rapid variability of seawater chemistry over the past 130 million years[END_REF]Yao et al., 2018), the severe restriction of the Mediterranean-Atlantic hydrologic exchanges that provoked the MSC profoundly disrupted the biogeochemical cycles of the Mediterranean sea, and specifically the short term carbon cycle (e.g., Damsté et al., 1995;Dela Pierre et al., 2012;Dela Pierre et al., 2014;Dela Pierre et al., 2015;Isaji et al., 2019;Natalicchio et al., 2019;[START_REF] Sabino | The response of water column and sedimentary environments to the advent of the Messinian salinity crisis: insights from an onshore deep-water section (Govone, NW Italy)[END_REF][START_REF] Sabino | Climatic and hydrologic variability in the northern Mediterranean across the onset of the Messinian salinity crisis[END_REF]. The effects of the MSC on the sulfur cycle, that is tightly coupled to the carbon cycle, however, have not been investigated to date.

The effect of Atlantic-Mediterranean restriction on the carbon cycle of the Mediterranean Sea is particularly visible in pre-MSC sediments, with: (i) the deposition of sapropel-diatomite alternations suggesting high primary productivity and carbon export, together with water column stratification (Pellegrino et al., 2018), (ii) the increasing number of independent proxies confirming bottom water stagnation and episodic anoxia to euxinia (Agiadi et al., 2022;Natalicchio et al., 2019;[START_REF] Sabino | The response of water column and sedimentary environments to the advent of the Messinian salinity crisis: insights from an onshore deep-water section (Govone, NW Italy)[END_REF][START_REF] Sabino | Climatic and hydrologic variability in the northern Mediterranean across the onset of the Messinian salinity crisis[END_REF]Vasiliev et al., 2019), and (iii) changes in foraminifer fauna from oxic to suboxic and overall drop in diversity (Bulian et al., 2022;Kouwenhoven et al., 1999;Kouwenhoven et al., 2003;[START_REF] Kouwenhoven | Paleoenvironmental evolution of the eastern Mediterranean during the Messinian: constraints from integrated microfossil data of the Pissouri Basin (Cyprus)[END_REF]Seidenkrantz et al., 2000).

By contrast to pre-MSC onset sediments, much less information on the biogeochemical functioning of Mediterranean environments during PLG gypsum deposition is available, since evaporites are generally devoid of biological proxies such as foraminifer fossils commonly used in paleoenvironmental reconstructions. Available evidence, however, suggest that Mediterranean environments could have hosted significant microbial communities, possibly fueling biogeochemical cycling of C and S (Dela Pierre et al., 2014;Isaji et al., 2019;Natalicchio et al., 2022;[START_REF] Natalicchio | Gypsum lithofacies and stratigraphic architecture of a Messinian marginal basin (Piedmont Basin, NW Italy)[END_REF]Schopf et al., 2012).

Early evidence comes from Vai and Ricci-Lucchi (1977) that assigned PLG gypsum to a microbial origin with the so-called "stromatolitic selenite" term, due to the visible occurrence of conspicuous filaments entrapped within gypsum. It is only recently however, that the microbial communities entrapped in PLG gypsum gained attention and have been identified as sulfur using consortia (Dela Pierre et al., 2015;[START_REF] Panieri | Microbial communities in Messinian evaporite deposits of the Vena del Gesso (northern Apennines, Italy)[END_REF][START_REF] Panieri | Ribosomal RNA gene fragments from G. Aloisi et al[END_REF]Schopf et al., 2012) and even more so that their widespread occurrence across the Mediterranean realm has been recognized (Natalicchio et al., 2022). The study of specific organic compounds (lipid biomarkers and geoporphyrins), as well as the study of diatoms, both particularly well preserved in PLG deposits, provided important information on the peculiar biogeochemical functioning of restricted basins during PLG formation (Isaji et al., 2019;Natalicchio et al., 2022;Pellegrino et al., 2021). Biogeochemical cycling during the formation of UG at the end of the MSC is even less documented. However, measurement of the oxygen isotopic composition of the sulfate making gypsum strongly suggests intense microbial cycling of sulfur during the formation of UG [START_REF] Aloisi | The geochemical riddle of "low-salinity gypsum" deposits[END_REF][START_REF] Waldeck | Calibrating the triple oxygen isotope composition of evaporite minerals as a proxy for marine sulfate[END_REF].

Although organic compounds and oxygen isotopic composition of sulfate analyses have been valuable to qualitatively assess the biogeochemical dynamics of Mediterranean basins during gypsum formation, they didn't allow to quantify the importance of these processes and the relative fluxes of elements involved. In this PhD work, I use the multiple sulfur isotope composition ( 34 S/ 32 S and 33 S/ 32 S) of gypsum and iron sulfide minerals to quantitatively investigate the biogeochemical cycling of sulfur during the deposition of gypsum contained in the Mediterranean Salt Giant.

Chapter 1 of this thesis consist in this general introduction that presents the main objects that I investigated in this work: the biogeochemical sulfur cycle as well as the Mediterranean salt giant. I also present the main tool that I used during my PhD: multiple sulfur isotopes.

In chapter 2 I present the two units I investigated during this work, the PLG and UG gypsum units that formed at the onset and the end of the MSC respectively. I also introduce the analytical techniques used in order to obtain the isotopic measurement that constituted the core of my approach and finally I present the numerical model and parameters that were used to interpret the results that will be presented in chapter 3.

In chapter 3 I present the results of an investigation of multiple sulfur isotopes in the Vena del Gesso marginal basin of the Mediterranean (northern Apennines, Italy). The aim is to: (i) quantify the relative importance of microbially driven reduction of sulfate and reoxidation of sulfide, compared to the burial of sulfate through gypsum precipitation (ii) quantify the impact of sulfur cycling on carbon cycling (iii) discuss and compare this ancient system to modern marine environments. The lead that convinced me to develop this study together with colleagues Francesco Dela Pierre and Marcello Natalicchio from Torino University and Pierre Cartigny (IPGP), comes from the occurrence of filaments of sulfide oxidizing bacteria and iron sulfide entrapped together within PLG gypsum which suggest the occurrence of both microbial sulfate reduction and sulfide reoxidation in marginal basins of the Mediterranean at that time (Dela Pierre et al., 2015;Natalicchio et al., 2022;Schopf et al., ). We combined petrographic observations with multiple sulfur isotope and numerical modelling to identify the biogeochemical processes and quantify the S and C fluxes involved during the formation of gypsum during the formation of the PLG unit in the Vena del Gesso basin (Northern Apennines, Italy). The results of this approach convinced me that the conspicuous filaments of sulfide oxidizing bacteria entrapped in gypsum had a crucial importance on the biogeochemistry of the Vena del Gesso basin. Large fluxes of sulfate were microbially reduced, with the fluxes of benthic carbon fueling this process being comparable to those of the most productive areas of the modern ocean where similar microbial consortia are found. In addition, the sulfide produced by MSR was almost quantitatively re-oxidized by the sulfide oxidizing bacteria. This allowed me to both (i) highlight the importance of sulfide reoxidative processes in the biogeochemistry of Messinian marginal basins of the Mediterranean and (ii) demonstrate that these processes can be identified and quantified using multiple sulfur isotope geochemistry and the careful understanding of the diagenetic environment of sulfate and sulfide bearing minerals. These results were published in Earth and Planetary Science Letters in June 2022 as "Cryptic sulfur cycling during the formation of giant gypsum deposits". An unformatted version of this scientific paper is presented as chapter 3 of this thesis including data tables and additional discussion within the supplementary informations at the end of the chapter.

I then decided to expand this approach to study the gypsum deposits that cover a larger area and were formed in the deep realm of the Mediterranean basin, providing a comparison between the biogeochemical cycling of sulfur in marginal and deep, as well as east and west, Mediterranean basins during the MSC. With this in mind, we decided to study the biogeochemical sulfur cycle of phase 3 gypsum, deposited in deep area of the Mediterranean (DSDP site 134,374,376 and ODP site 654), that lie on top of the phase 2 halite layer and below the Miocene-Pliocene boundary. The results of this approach are presented in chapters 4 and 5 of this manuscript.

In chapter 4, I collaborated with ESR2 Hanneke Heida (CESIC, Barcelona, Spain) and ESR5

Federico Andreetto (Utrecht University, Netherlands) of the SaltGiant ETN to investigate some of the most controversial issues of the MSC: (i) What was the depth and water column thicknesses that characterized Mediterranean basins during UG formation, after the acme of the MSC and before the reestablishment of normal marine conditions in the Mediterranean? This allowed us to tackle the question: (ii) What was the minimal magnitude of the Mediterranean level drop during the MSC? To do so, I investigated the effect of water column thickness on the δ 34 S isotopic composition of sedimentary sulfide. This allowed to obtain an estimation of the maximal water column thickness under which the iron sulfide we investigated was formed. We then used this estimate as an input for the paleo-bathymetric reconstruction of each individual basin investigated. Once the paleo-depth of each basin was obtained, it was utilized to estimate the minimal amplitude of the level fall in both western and eastern Mediterranean basins. To match my isotopic observations, a level fall of at least 1.1 km in the western Mediterranean basin and of at least 2.5km in the eastern Mediterranean basin has to be considered, suggesting that these two basins have been disconnected at least intermittently during the MSC. This study provides for the first time a biogeochemical evidence for a large drawdown of the Mediterranean level and the disconnection of the two main basins of the Mediterranean.

In chapter 5, I present an isotopic investigation of the same iron sulfide samples studied in chapter 4 but considering also the minor sulfur isotope ratio 33 S/ 32 S. I also included isotopic measurement of the sulfate molecule within gypsum ( 33 S/ 32 S, 34 S/ 32 S and 18 O/ 16 O). This approach was used with simple numerical modelling of the effect of Rayleigh distillation, mixing and various redox processes effect to finely reconstruct the dynamics of the biogeochemical cycling of sulfur during the formation of UG gypsum. Specifically, the goal was to: (i) discriminate the effects of sulfide pooling from changes in microbial sulfate reduction rates on the multiple sulfur isotope composition of iron sulfide during UG formation (ii) quantify the relative fluxes of sulfur that were involved in sulfide reoxidation. The results of this study allowed us to confirm the occurrence of a very active biogeochemical cycling of sulfur in deep basins that cover a large east-west transect of the Mediterranean, with an almost quantitative reoxidation of the sulfur reduced by MSR. While we already observed a similar "cryptic" cycling of sulfur in a small marginal basin of the Mediterranean at the onset of the MSC (chapter 3), this confirmed that this biogeochemical functioning was common during the MSC in open, deep Mediterranean basins and not only restrained to small, marginal and shallow water basins.

In chapter 6 I present the general conclusions of this work and its perspectives for investigating the sulfur cycle during not only the phenomenal MSC event, but also in other past or present open marine environments.

In the appendix section, I present the dataset used in this study as well as the scientific papers The oxidation and reduction of sulfur compounds play a fundamental role in Earth's elemental cycles. In many environments, microbial reduction of sulfate into sulfide largely fuels the anaerobic oxidation of carbon. Under some conditions, microbially produced sulfide can be completely re-oxidized back to sulfate. This process is obscured in sulfur concentration profiles, leading to an underestimation of microbial sulfate reduction rates because the observed net sulfate consumption (and sulfide production) is zero. Complete re-oxidation of sulfide is also hidden in most measurements of δ 34 S and δ 18 O isotopes, making this cryptic sulfur cycling challenging to track in modern environments, and even harder to detect in geological samples representing past environments where pore water analyses are not available.

The determination of multiple nontraditional sulfur isotopes (Δ 33 S and Δ 36 S) enables sulfide re-oxidative processes to be distinguished from concomitant microbial involving sulfur produce a small deviation from the expected mass-dependent relationship between 33 S, 34 S and 36 S. The determination of multiple sulfur isotopes first requires that the S species of interest (often sulfate and sulfide) is chemically extracted and transformed to pure SF 6 by a number of intermediate steps. The sulfur isotopes are then analyzed by gaseous mass spectrometry. In the δ 34 S-Δ 33 S space, the relative data points position will differ if the measured sulfur is the result of a mixing between different sources of sulfur or of redox processes that fractionate sulfur isotopes. By using fractionation factors determined in laboratory experiments for a range of redox processes in biogeochemical models of sulfur, it is possible to disentangle complex sulfur cycling where several processes and mixing might occur altogether. Highly precise measurements are possible in samples containing as little as 0.1µmol of sulfur, and thereby allow the characterization of both major and minor sulfur phases of an environmental sample.

The simultaneous measurement of δ 34 S, Δ 33 S and Δ 36 S has revealed sulfide oxidation hypersaline cold springs and sulfide-rich cave systems. This tool thus can help refine rates of organic carbon consumption by microbial sulfate reduction, which are potentially underestimated in environments where re-oxidative sulfur cycling is particularly prominent. This tech nique is also useful to unravel past cryptic sulfur cycling, as recorded in gypsum from the late Miocene, for example, and might enable a deeper understanding of complex redox cycling processes in modern environments that are sensitive to anthropogenic effects. signature of biogenic and abiotic sulfur redox processes, multiple sulfur isotope analysis of extra-terrestrially derived sulfur-rich mineral samples could provide evidence of life on other planets.

Stable Isotopes Geochemistry Lab, Institut de Physique du

Globe de Paris, Université de Paris, Paris, France.

e-mail: guibourdenche@ipgp.fr 1.8.2 Stable isotopes of sulfur: how and why they are useful to decrypt sulfur cycling A majority of chemical elements have several isotopes, i.e. they have the same number of protons (that define the nature of the element) but a variable number of neutrons. The isotopes of a same chemical element have thus slightly different masses but similar chemical properties. In nature, sulfur has 4 stable isotopes with different abundances and masses , 32 S (mass 32) representing 95.02% of the total sulfur, 33 S (mass 33) (0.76%), 34 S (mass 34) (4.22%) and 36 S(mass 34) (0.01%) (Beaudoin et al., 1994;[START_REF] Canfield | Biogeochemistry of sulfur isotopes[END_REF]Kaplan and Hulston, 1966). The isotopic composition of a given sulfur isotope in a sample is expressed as the ratio between the less abundant isotope ( 33 S, 34 S or 36 S) and the most abundant isotope ( 32 S) measured for this sample normalized by the same ratio measured for a standard, generally international Canyon Diablo Troilite (CDT) or Vienna-CDT (synthesized) for sulfur isotopes [START_REF] Brand | Assessment of international reference materials for isotope-ratio analysis (IUPAC Technical Report)[END_REF]Kaplan and Hulston, 1966). This deviation (noted δ) is expressed in ‰.

𝛿 "# 𝑆 = ⎝ ⎛ "# 𝑆 ") 𝑆 *+,-./ "# 𝑆 ") 𝑆 0123 -1

⎠ ⎞ × 1000

The differences in masses between isotopes implies differences in the reactions in which they are involved, the lighter isotopes generally reacting more rapidly than the heavier. The differences of mass. These differences in behavior often result in the different partitioning of the various isotope of an element while used in a reaction and is expressed by fractionation factors α (Dauphas and Schauble, 2016).

𝛼 ;→= = > "# 𝑆 ") 𝑆 ? = > "# 𝑆 ") 𝑆 ?

;

With α the fractionation factor during the reaction of sulfur from A to B and 3x S/ 32 S the isotopic ratios of the reactant A and B the product with x=34 or 33.

For simplicity sakes, this fractionation factor can also expressed in ‰ and in that case become ε.

𝜀 ;→= = (𝛼 ;→= -1) × 1000

Because sulfur has more than two stable isotopes, when considering both 34 S/ 32 S and 33 S/ 32 S isotopic ratios, mass-dependent fractionating processes follow a high temperature approximation when the reaction involved is at equilibrium (Young et al., 2002). This approximation follows a curve in the δ 34 Sδ 33 S space:

C 𝛿 "" 𝑆 1000 + 1E = C 𝛿 "F 𝑆 1000 + 1E G/H
Where the slope 𝜆 /H = J

K J L J K M J K J L J K N
And m1<m2<m3 corresponding to the masses of the isotopes considered. Here m1 corresponds to 32 S, m2 to 33 S and m3 to 34 S. At equilibrium, λnet =0.515 ( Figure 1. 18).

If a process occurs at thermodynamic equilibrium, both the product and the reactants will fall on the equilibrium fractionation curve and result in a ∆ 33 S = 0, with ∆ 33 S representing the deviation from the equilibrium fractionation curve. When kinetic processes are at play, a significant deviation from this _________________________________________________________________________________________________ 40 equilibrium fractionation curve can be detected thanks to recent advances in mass spectrometric measurements (Johnston, 2011).

Figure 1. 18 Diagram showing the partitioning of isotopes along the mass dependent fractionation line at equilibrium

In such case, the partitioning of sulfur isotope sometimes follows different fractionation laws, that are often processes specific. Because the Δ 33 S notation describes the difference between the equilibrium curve (δ 34 Seq-δ 33 Seq) and the curve on which would lie the products of a given process (δ 34 Sprocess -δ 33 Sprocess), ∆ 33 S ≠ 0 if the considered process is out of equilibrium (Figure 1. 19). ∆ 33 S>0 when the fractionation curve lies above the equilibrium array and ∆ 33 S<0 when it lies below the equilibrium array. Thus, in the δ 34 -∆ 33 S space, product of processes following different mass fractionation laws than equilibrium will follow a straight line (Ono et al., 2006) (see Figure 1. 19). correspond to the two poles of the mixing in the δ 34 S-δ 33 S space. The red line represents the line on which the product of the mixing fall in the δ 34 S-δ 33 S space, its relative position the the red line depending on the relative proportion of each endmember in the final product. In δ 34 S-∆ 33 S space, the product of the mixing fall on the solid black curve.

Non-zero ∆ 33 S can also arise from mixing processes. Mixing of two end-members with distinct isotopic compositions (δ 33 S1, δ 33 S2 and δ 34 S1, δ 34 S2) produces a distinct isotopic reservoir with a third isotopic composition (δ 33 Smix, δ 34 Smix) that depends on the isotopic composition of the end members and their relative proportion in the mix (f1 and f2=1-f1).

This relationship is written as follows:

δ 34 Smix= δ 34 S1f1+ δ 34 S2(1-f1) δ 33 Smix= δ 33 S1f1+ δ 33 S2 (1-f1) In the δ 33 S-δ 34 S space, this mixing describes a line (Figure 1. 20). The mixing is thus the difference between the fractionation curve and the mixing line: Δ 33 Smix= δ 33 Smix-δ 33 Seq. In the Δ 33 S-δ 34 S space, this mixing thus follows a curve: as the distance between the equilibrium curve and the mixing line increase, Δ 33 S increases. Whereas when the mixing line crosses the equilibrium curve, Δ 33 S=0 (Figure 1. 20).

In terrestrial environments such as those studied in this work, the sulfur isotope composition of natural materials evolves as sulfur pools are mixed and processed. The use of δ 34 S-∆ 33 S diagrams can help to obtain a first order idea of the different sulfur pools and processes involved in the cycling of sulfur in a given system. In addition, fractionation factors 34 Ɛnet and 33 λ net, calculated between co-eval sulfate and sulfide, are insensitive to mixing and do not depend on the starting isotopic composition of reacting sulfur species, but rather on the processes at play during sulfur cycling. Laboratory experiments with sulfur-metabolizing microbes have defined ranges of 34 Ɛ-33 λ values for a variety of biogeochemical processes (Eldridge and Farquhar, 2018;Johnston et al., 2005;[START_REF] Zerkle | Fractionation of multiple sulfur isotopes during phototrophic oxidation of sulfide and elemental sulfur by a green sulfur bacterium[END_REF].

These ranges can be compared with the 34 Ɛnet and 33 λnet values obtained by measurements of sulfate and pyrite co-occurring in natural samples in order to obtain a first-order assessment of the biogeochemical processes at play during their genesis. (Eldridge and Farquhar, 2018) (dark grey area with black dotted contour), data from pure cultures with sulfate reducers comes from (Bertran et al., 2018;Johnston et al., 2005;Leavitt et al., 2013;Sim et al., 2011a;and Sim et al., 2011b) (light grey area) and data from pure culture experiments with disproportionators comes from (Johnston et al., 2005) (white area with grey dashed contour). White squares represent the 34 ɛnet and 33 λnet measured in natural environments populated with biofilms dominated by chemolithotroph sulfide oxidizer (Magnuson et al., 2020;Zerkle et al., 2016).

Chapter 2. Lofi (2018) The main work that I undertook during my PhD focuses on the geochemistry of gypsum formed during the MSC. There are numerous gypsum deposits nowadays founds in the Mediterranean area that correspond to Messinian gypsum (Figure 2. 1). They consist of Primary Lower Gypsum (PLG) deposits that formed during the first phase of the MSC, resedimented PLG, called Resedimented Lower Gypsum (i.e. RLG) that is the result of the It includes gypsum initially deposited in peripheral basins of the Mediterranean and nowadays exposed on land and/or mapped offshore and accessed through seismic imaging and/or borehole dataset (see the green areas and the orange dots in Figure 2. 1). Gypsum deposits formed during the MSC are found all around the Mediterranean area : in Albania (Pashko et al., 2019), Algeria [START_REF] Rouchy | The end of the Messinian salinity crisis: evidences from the Chelif Basin (Algeria)[END_REF], Adriatic foredeep (Manzi et al., 2020), Balearic Promontory (Ochoa et al., 2015;Raad et al., 2021), Calabria, Piedmont, Northern Apennines, Tuscany or Sicily in Italy [START_REF] Costanzo | Messinian twinned selenite from the Catanzaro Trough, Calabria, Southern Italy: field, petrographic and fluid inclusion perspectives[END_REF]Dela Pierre et al., 2011;Dinelli et al., 1999;Lugli et al., 2007;Lugli et al., 2010;[START_REF] Natalicchio | Did Late Miocene (Messinian) gypsum precipitate from evaporated marine brines? Insights from the Piedmont Basin (Italy)[END_REF][START_REF] Natalicchio | Gypsum lithofacies and stratigraphic architecture of a Messinian marginal basin (Piedmont Basin, NW Italy)[END_REF], Southern Spain (Lu et al., 2001;Lugli et al., 2010), Cyprus [START_REF] Robertson | Depositional processes and basin analysis of Messinian evaporites in Cyprus[END_REF][START_REF] Rouchy | Stromatolites and Cryptalgal Laminites Associated with Messinian Gypsum of Cyprus[END_REF], Egypt (Aref, 2003;[START_REF] Melegy | Microstructure and geochemistry studies on Messinian gypsum deposits from the Northern Coast of Egypt[END_REF], Israel (Cohen, 1993;Raab et al., 1997), Tunisia (El Euch-El Koundi et al., 2009;Harrab et al., 2013;Sghari, 2012), Turkey (Boulton et al., 2016;Tekin et al., 2010) and Syria (Hardenberg and Robertson, 2007;[START_REF] Rouchy | The end of the Messinian salinity crisis: evidences from the Chelif Basin (Algeria)[END_REF] or gypsum drilled during DSDP/IODP oceanographic expeditions LEG 13,42 and 107 (see the green dots in Figure 2. 1) (Hsü et al., 1978a;[START_REF] Hsü | Initial reports of the deep sea drilling project: DSDP volume XLII Part[END_REF]Hsü et al., 1978d;Hsü et al., 1973;[START_REF] Kastens | Proc. ODP, Init. Repts[END_REF][START_REF] Lugli | The deep record of the Messinian salinity crisis: evidence of a non-desiccated Mediterranean Sea[END_REF].

Geological Background

In this thesis we've been mostly interested in PLG and UG deposits formed respectively at the onset and at the end of the MSC. In the following sections, we present background information on these deposits and on specific sections and sites that we've investigated.

Chapter 2-Material and Methods ___________________________________________________________________________________________ 47 2.1 Geological settings of Primary Lower Gypsum (PLG) deposits

General Geological background

The Primary Lower Gypsum (PLG) unit was deposited during the first stage of the Messinian Salinity Crisis (5.97-5.60 Ma (Manzi et al., 2013)) in shallow (<200m) to intermediate basins (<800) of the Mediterranean area, including Southern Spain, Italy, Greece and Cyprus (Karakitsios et al., 2017;Krijgsman et al., 1999;Lugli et al., 2010) (Figure 2. 2).

Figure 2. 2 Distribution of Messinian Evaporites in the Mediterranean and location of the Vena del

Gesso basin (pink star). Map modified after Lofi (2018) This unit was deposited in approximately 380 kyrs (Roveri et al., 2014a) and is formed by lithological cycles consisting of marl (and carbonate beds)/gypsum couplets (Figure 2. 3).

Deposition of these cycles was controlled by precession-driven climate oscillations (Krijgsman et al., 1999;Lugli et al., 2010;Topper et al., 2014). Tuning of the lithological cycles to the Chapter 2-Material and Methods ___________________________________________________________________________________________ 48 astronomical reference curve [START_REF] Laskar | La2010: a new orbital solution for the long-term motion of the Earth[END_REF] allows the precise dating of each cycle and bed by bed correlation of the PLG unit at the Mediterranean scale (see the pink dots in Figure 2. 2) (Karakitsios et al., 2017;Krijgsman et al., 1999;Lugli et al., 2010;[START_REF] Orszag-Sperber | The onset of the Messinian salinity crisis: insights from Cyprus sections[END_REF]. Large sections of PLG deposits can be found nowadays outcropping on land (Figure 2.

2). This is due to tectonic activity during/or after the MSC that allowed the exhumation of shallow marginal basin deposits onshore and have favored their characterization. PLG deposits in Mediterranean marginal basins (Lugli et al., 2010) have particular features that are observed everywhere across the Mediterranean realm.

This includes:

(i) the presence of similar facies and stacking patterns in marginal basins thousands of kilometers apart suggesting the concomitant and subaqueous deposition of PLG at the Mediterranean scale [START_REF] Ciesm | The Messinian Salinity Crisis from mega deposits to microbiology[END_REF]Lugli et al., 2010) (Figure 2.

3)

Chapter 2-Material and Methods ___________________________________________________________________________________________ 49 (ii) the alternation of marl or carbonate beds with gypsum deposits correlated at the basin scale that are the lithological expression of cyclical changes in basin chemistry driven by precessionally-forced climate oscillations [START_REF] Ciesm | The Messinian Salinity Crisis from mega deposits to microbiology[END_REF]Flecker et al., 2015;Karakitsios et al., 2017;Krijgsman et al., 1999;Lugli et al., 2007;Lugli et al., 2010;Topper et al., 2011) (Figure 2. 3) (iii) the occurrence of marine biota together with strontium and sulfate isotopic composition of gypsum close to Messinian oceanic values (Carnevale et al., 2019;García-Veigas et al., 2018;Lu et al., 2001;Lugli et al., 2007;Lugli et al., 2010;Pellegrino et al., 2021;[START_REF] Reghizzi | Isotope stratigraphy (87Sr/86Sr, δ18O, δ13C) of the Sorbas basin (Betic Cordillera, Spain): Paleoceanographic evolution across the onset of the Messinian salinity crisis[END_REF]Reghizzi et al., 2018;Roveri et al., 2014a) (e.g. Figure 2. 4). This indicates that PLG deposits formed in basins still connected to the Mediterranean Sea [START_REF] Aloisi | The geochemical riddle of "low-salinity gypsum" deposits[END_REF]Flecker et al., 2015;Topper et al., 2014;Topper et al., 2011).the lateral transition of gypsum and marls alternation in marginal basins into organic rich sediments and carbonates in deeper settings [START_REF] Ciesm | The Messinian Salinity Crisis from mega deposits to microbiology[END_REF]Dela Pierre et al., 2011;[START_REF] Manzi | The onset of the Messinian salinity crisis in the deep Eastern Mediterranean basin[END_REF][START_REF] Gennari | A shallow water record of the onset of the Messinian salinity crisis in the Adriatic foredeep (Legnagnone section, Northern Apennines)[END_REF][START_REF] Manzi | The deep-water counterpart of the messinian lower evaporites in the apennine foredeep: the fanantello section (northern Apennines, Italy)[END_REF]Roveri et al., 2008;[START_REF] Roveri | The Messinian salinity crisis: Looking for a new paradigm? Palaeogeography[END_REF]. Some authors suggest that the inhibition of gypsum precipitation in the deep part of the basin is controlled by the consumption of sulfate at the bottom of the basin (de Lange and Krijgsman, 2010;Dela Pierre et al., 2012;Natalicchio et al., 2019) (Figure 2.5). Lugli et al. (2010).

Chapter 2-Material and Methods ___________________________________________________________________________________________ 51 (iv) the preservation of microbial remains entrapped in gypsum in Cyprus (Allwood et al., 2013;Natalicchio et al., 2022;[START_REF] Rouchy | Gypsum Microbial Sediments: Neogene and Modern Examples[END_REF][START_REF] Rouchy | Stromatolites and Cryptalgal Laminites Associated with Messinian Gypsum of Cyprus[END_REF], Crete (Allwood et al., 2013;Natalicchio et al., 2022;[START_REF] Rouchy | Gypsum Microbial Sediments: Neogene and Modern Examples[END_REF], Ionian Islands [START_REF] Rouchy | Gypsum Microbial Sediments: Neogene and Modern Examples[END_REF], Algeria [START_REF] Rouchy | Gypsum Microbial Sediments: Neogene and Modern Examples[END_REF], Southern Spain (Natalicchio et al., 2022), Northern Apennines (Natalicchio et al., 2022;[START_REF] Panieri | Microbial communities in Messinian evaporite deposits of the Vena del Gesso (northern Apennines, Italy)[END_REF][START_REF] Panieri | Ribosomal RNA gene fragments from G. Aloisi et al[END_REF][START_REF] Rouchy | Gypsum Microbial Sediments: Neogene and Modern Examples[END_REF]Vai and Ricci-Lucchi, 1977), Piedmont (Dela Pierre et al., 2015) and Sicily (Hardie and Eugster, 1971;[START_REF] Rouchy | Gypsum Microbial Sediments: Neogene and Modern Examples[END_REF] (Figure 2. 6).

These so-called « spaghetti-like » microbial fossils in coeval gypsum deposits thousands of km apart are interpreted by authors to be the remnants of cyanobacteria [START_REF] Panieri | Microbial communities in Messinian evaporite deposits of the Vena del Gesso (northern Apennines, Italy)[END_REF][START_REF] Panieri | Ribosomal RNA gene fragments from G. Aloisi et al[END_REF][START_REF] Rouchy | Gypsum Microbial Sediments: Neogene and Modern Examples[END_REF] or sulfide oxidizing bacteria (Carnevale et al., 2019;Dela Pierre et al., 2015;Schopf et al., 2012) and reflect anyway particular biogeochemical settings during the formation of gypsum (Figure 2. 6). These studies suggest different paleoenvironmental meanings, with possible presence of microbial sulfate reduction activity (Dela Pierre et al., 2015;Natalicchio et al., 2022;[START_REF] Panieri | Microbial communities in Messinian evaporite deposits of the Vena del Gesso (northern Apennines, Italy)[END_REF]Schopf et al., 2012) either in shallow hypersaline restricted basins [START_REF] Panieri | Microbial communities in Messinian evaporite deposits of the Vena del Gesso (northern Apennines, Italy)[END_REF][START_REF] Panieri | Ribosomal RNA gene fragments from G. Aloisi et al[END_REF][START_REF] Rouchy | Gypsum Microbial Sediments: Neogene and Modern Examples[END_REF] or in a stratified water column with anoxic bottom waters (Dela Pierre et al., 2015;Schopf et al., 2012).

In this study we have chosen to investigate the Vena del Gesso section, from which the earliest evidence of microbial filaments have been identified. We describe the particular settings of this basin in the following section. Mediterranean Sea (Lugli et al., 2007;[START_REF] Manzi | Deep-water clastic evaporites deposition in the Messinian Adriatic foredeep (northern Apennines, Italy): did the Mediterranean ever dry out?[END_REF][START_REF] Roveri | Sedimentary and tectonic evolution of the Vena del Gesso basin (Northern Apennines, Italy): Implications for the onset of the Messinian salinity crisis[END_REF]. [START_REF] Laskar | La2010: a new orbital solution for the long-term motion of the Earth[END_REF] Chapter 2-Material and Methods

___________________________________________________________________________________________

55

We sampled gypsum and marls from PLG cycle 1-2 in the Monticino quarry (44°130 30ʺN, 11°450 44ʺE) (Marabini and Poluzzi, 1977) (Figure 2. 9) whereas cycles 3 to 16 were sampled in the Monte Tondo quarry (44°150 8ʺN, 11°400 6ʺE) (Lugli et al., 2010;Manzi et al., 2013) (Figure 2. 10). anoxia [START_REF] Carnevale | Did the Mediterranean marine reflooding precede the Miocene/Pliocene boundary? Paleontological and geochemical evidence from upper Messinian sequences of Tuscany, Italy. Paleogeogr[END_REF]. The molecular fossil inventory indicates that the water column was probably stratified and periodically euxinic up to the euphotic zone, and that the organic matter deposited on the basin floor underwent rapid and extensive sulfurization (Damsté et al., 1995;Isaji et al., 2019;Keely et al., 1995;[START_REF] Kenig | Molecular Indicators for Paleoenvironmental Change in a Messinian Evaporitic Sequence (vena Del Gesso, Italy). 2. High-Resolution Variations in Abundances and C-13 Contents of Free and Sulfur-bound Carbon Skeletons in a Single Marl Bed[END_REF]Schaeffer et al., 1995). The thin marl beds are followed of the massive selenite lithofacies (Lugli et al., 2010) (Figure 2. 12). The absence of dissolution features suggest that this facies is a fully subaqueous deposit formed in a relatively deep basin permanently covered by supersaturated brines (Lugli et al., 2010). The selenite crystals show an internal lamination given by the alternation of mm-thick turbid and limpid laminae, possibly reflecting seasonal climate variability [START_REF] Orti | Selenite facies in marine evaporites: A review[END_REF]Dela Pierre et al., 2015;Reghizzi et al., 2018) (Figure 2. 13 Figure 2. 14). In the limpid laminae, solid inclusions are scarce or even absent. In contrast, the turbid laminae entrap densely interwoven filamentous fossils (Lugli et al., 2010;[START_REF] Panieri | Ribosomal RNA gene fragments from G. Aloisi et al[END_REF]Schopf et al., 2012;Natalicchio et al., 2021) that are observed both in the re-entrant angles of the twins and along the vertical growth bands (Figure 2. 6Figure 2. 13). Clay-rich aggregates (marine snow floccules) and nanosized planktic diatoms are also observed in the turbid laminae (Pellegrino et al., 2021) (Figure 2. 15). Starting from the gypsum beds of the 4th PLG cycle, the massive selenite facies is followed upwards by the banded selenite facies composed of less than 10 cm-thick crusts of bottom grown twinned crystals separated by thin carbonate laminae (Lugli et al., 2010) (Figure 2. 8 Figure 2. 16). This facies possibly reflects cyclical vertical fluctuation of the gypsum saturation interface (i.e the pycnocline) (Babel, 2007;Lugli et al., 2010) (Figure 2. 16). Filamentous microfossils, similar to those found in the massive selenite, as well as dolomite and pyrite have been described in the banded selenite of the 6 th cycle of the Vena del Gesso basin [START_REF] Panieri | Microbial communities in Messinian evaporite deposits of the Vena del Gesso (northern Apennines, Italy)[END_REF].From the 6th PLG cycle to the top of the section, massive and banded selenite are followed by the branching selenite facies consisting of cm-sized clear selenite crystals growing at the seafloor inclined or horizontal and grouped together in nodules or lenses separated by thin fine grained carbonate or gypsum laminae (Lugli et al., 2010;[START_REF] Natalicchio | Gypsum lithofacies and stratigraphic architecture of a Messinian marginal basin (Piedmont Basin, NW Italy)[END_REF] (Figure 2. 17). We analyzed a total of 88 samples from the Vena del Gesso basin. Samples were collected for the different representative lithofacies of each cycle. In gypsum, we analyzed the multiple sulfur and oxygen isotope composition of the sulfate ion (δ 34 SSO4, Δ 33 SSO4 and δ 18 OSO4, respectively) (see section 1.9 for introduction on stable isotope systems). The multiple sulfur isotope composition of iron sulfides (δ 34 SIS and Δ 33 SIS) was also analyzed in both gypsum and marly interbeds (see section 2.3.1 and 2.3.2 for details on the isotopic measurements).

Sulfides were chemically extracted from the gypsum layers (of which they are a very minor component) and the marls (see section 2. The Upper Gypsum (UG) unit was deposited during the third stage of the MSC (5.55-5.332 Ma in deep (>800m depth) to intermediate (200 to 800m depth) basins of the Mediterranean (Andreetto et al., 2021a). Similarly to PLG deposits, the UG consists in a cyclic alternation of calcium sulfate and marls/mudstone interbeds that are interpreted to be controlled by precession driven climate oscillations [START_REF] Manzi | A new facies model for the Upper Gypsum of Sicily (Italy): chronological and paleoenvironmental constraits for the Messinian salinity crisis in the Mediterranean[END_REF]Roveri et al., 2014b) (Figure 2. 20).

The paleontological and sedimentological expression of phase 3 deposits is complex and vary amongst different geological settings, making the significance of this record in terms of paleoenvironment largely debated (Andreetto et al., 2021a). In the intermediate depth basins of Sicily and Cyprus, UG outcrops on land and the gypsum is found mainly in the form of both bottom-grown selenite, cumulate gypsum (and to a smaller extent clastic gypsum (Andreetto et al., 2022a;[START_REF] Manzi | A new facies model for the Upper Gypsum of Sicily (Italy): chronological and paleoenvironmental constraits for the Messinian salinity crisis in the Mediterranean[END_REF]Manzi et al., 2016) (García-Veigas et al., 2018;Grossi et al., 2015;Manzi et al., 2016;[START_REF] Rouchy | Paleoenvironmental changes at the Messinian-Pliocene boundary in the eastern Mediterranean (southern Cyprus basins): significance of the Messinian Lago-Mare[END_REF]. In the non-evaporitic interbeds of the same sections, shallow water brackish ostracods of Paratethyan affinity are found and display lower strontium isotopic values than the co-eval Messinian ocean (Figure 2. 21 and Figure 2. 22). This has been interpreted to reflect the dilution of the Mediterranean sub-basins with continental waters during marl deposition (García-Veigas et al., 2018;Grossi et al., 2015;Manzi et al., 2016;[START_REF] Rouchy | Paleoenvironmental changes at the Messinian-Pliocene boundary in the eastern Mediterranean (southern Cyprus basins): significance of the Messinian Lago-Mare[END_REF].

In Cyprus, phase 3 deposits present erosional unconformities or paleosols, indicative of the occasional subaerial exposure of some marginal basins of the Mediterranean during phase 3 [START_REF] Orszag-Sperber | The onset of the Messinian salinity crisis: insights from Cyprus sections[END_REF][START_REF] Rouchy | Paleoenvironmental changes at the Messinian-Pliocene boundary in the eastern Mediterranean (southern Cyprus basins): significance of the Messinian Lago-Mare[END_REF]. Both these evidences suggest at least intermittent disconnection of these basins from marine influxes during phase 3. Conversely, marine organisms have also been found in non-evaporitic interbeds of these basins and

Chapter 2-Material and Methods ___________________________________________________________________________________________ 68 although their pristine origin is controversial, some authors suggest that their widespread occurrence in marginal to deep basins of the Mediterranean indicates that fully marine conditions were maintained throughout phase 3 (Carnevale et al., 2019;Carnevale and Schwarzhans, 2022). Finally, in more marginal areas where basins are thought to have been shallower than Caltanissetta (Sicily) and southern Cyprus basins, phase 3 deposits are devoid of evaporites.

They are mostly constituted by terrigenous deposits containing Paratethyan brackish water fauna similar to the above-mentioned intermediate basins, with strontium isotopic composition deviating from Messinian sea water value (Andreetto et al., 2021a;Andreetto et al., 2022b;Andreetto et al., 2021b) (see blue, grey and red dots in Figure 2. The scenarios for phase 3 are largely based on sedimentary successions nowadays outcropping on land. However, the bulk of UG deposits are nowadays buried under the sea floor and were mainly studied through seismic studies (Lofi, 2018;Lofi et al., 2011a;Roveri et al., 2014a). These deposits are thick and cover most of the area from the Alboran volcanic arc

22). Although it is

Chapter 2-Material and Methods ___________________________________________________________________________________________ 70 to the Ionian basin, whereas they are absent or only 100m to 300m thick in some areas of the Eastern Mediterranean [START_REF] Gvirtzman | Intra-Messinian truncation surface in the Levant Basin explained by subaqueous dissolution[END_REF]Lofi, 2018) (see green areas in Figure 2. 19 and ). They are generally concentrated in depressions and abyssal plains, and they pinch out at the foot of continental slopes and seamounts (Lofi et al., 2011b). Few IODP and DSDP expeditions successfully drilled the uppermost part of phase 3 Messinian deposits, confirming the occurrence of gypsum on top of phase 2 halite massive deposits in a variety of water depths across the Mediterranean, (see green dots and stars in Figure 2. 19) [START_REF] Comas | Proceedings of the Ocean Drilling Program, Initial Reports[END_REF]Hsü et al., 1973;[START_REF] Kastens | The geological evolution of the Tyrrhenian Sea: An introduction to the scientific results of ODP Leg 107[END_REF]. The relative position of gypsum-marls alternation on top of halite deposits and below the Miocene-Pliocene boundary assign the gypsum drilled during ODP/DSDP expeditions to the final phase of the MSC (Hardie and Lowenstein, 2004;[START_REF] Lugli | The deep record of the Messinian salinity crisis: evidence of a non-desiccated Mediterranean Sea[END_REF]Roveri et al., 2014b).

However, the poor recovery obtained during the drilling campaigns almost never allowed for a direct correlation between the cores nor with onshore UG deposits making a detailed dating of UG deposits sampled in IODP and DSDP cores impossible (Roveri et al., 2014b). The gypsum from these cores consists of cumulate, clastic or porphyroblastic gypsum as well as anhydrite (Hardie and Lowenstein, 2004;[START_REF] Lugli | The deep record of the Messinian salinity crisis: evidence of a non-desiccated Mediterranean Sea[END_REF] Cumulate gypsum corresponds to mm-thick rhythmic successions of gypsum laminae (Figure 2. 26). The individual crystals were formed at the brine air interface or within the water column and settled to the bottom of the basin by gravity (meaning either falling all the way the bottom of the basin or gliding along the slopes). Sometimes, primary cumulate layers

___________________________________________________________________________________________

73 alternate with fine grained clastic layers originating from the reworking of cumulate sediments by low density gravity flows [START_REF] Lugli | The deep record of the Messinian salinity crisis: evidence of a non-desiccated Mediterranean Sea[END_REF]. The non-evaporitic interbeds in deep basins, similarly to their more marginal counterparts, contain various fossils assemblages that have led to controversial interpretations of the depositional environments under which they accumulated (see Andreetto et al., 2021a in Appendix D for an extensive review of the debate).

On one hand, planktonic foraminifera and monospecific nannofossils assemblages point to Mediterranean basins connected to oceanic waters during phase 3 marls interbeds deposition (e.g [START_REF] Popescu | Lago Mare episodes around the Messinian-Zanclean boundary in the deep southwestern Mediterranean[END_REF]. On the other hand, benthic fossils assemblages suggest the occurrence of brackish, shallow water to neritic environments in abyssal plains of the Mediterranean suggesting that the Mediterranean water level was so low that its various basins were disconnected from each other (Hsu et al., 1978;Hsü et al., 1973;[START_REF] Reiche | The impact of salt on the late Messinian to recent tectonostratigraphic evolution of the Cyprus subduction zone[END_REF]Ryan et al., 1973). Whether or not some or all of these fossil assemblages are in-place white and gray millimetric laminae (Ryan et al., 1973) (Figure 2. 31). White laminae are made of pure anhydrite, while the gray bands are made of mud (Ryan et al., 1973). This anhydrite layer is found in contact with halite. A similar but thicker sequence is present in the last core (core 10) of the well.

Here the halite bed (Section 2 and 3 and core catcher) is 1.6 m thick below a 40 cm-thick laminar anhydrite with nodules [START_REF] Lugli | The deep record of the Messinian salinity crisis: evidence of a non-desiccated Mediterranean Sea[END_REF] and 30 cm of dark gray, hydrocarbon-rich dolomitic marls including rare planktonic foraminifera nannoplankton together with ostracods, radiolarian debris and sponge spicules (Figure 2. 31). The halite bed is composed of thin-bedded accumulations of fine-grained halite cube crystals (laminar halite facies of [START_REF] Lugli | The deep record of the Messinian salinity crisis: evidence of a non-desiccated Mediterranean Sea[END_REF]. This intra-UG salt layer was first interpreted as autochthonous and deposited in a sabkha established as a consequence of an important phase of base-level lowering [START_REF] Geletti | The Messinian Salinity Crisis: new seismic evidence in the West-Sardinian Margin and Eastern Sardo-Provençal basin (West Mediterranean Sea)[END_REF]Hsu et al., 1978;[START_REF] Mocnik | The Messinian Salinity Crisis in the West-Mediterranean Basins: comparison between two rifted margins[END_REF]. However, other authors suggest that the lack of evidence of desiccation (Hardie and Lowenstein, 2004;[START_REF] Lugli | The deep record of the Messinian salinity crisis: evidence of a non-desiccated Mediterranean Sea[END_REF] and the presence of salt diapirism indicate that this halite deposited during phase 2 and flowed upward during salt tectonic deformation processes (Andreetto et al., 2021a et al., 1989;[START_REF] Borsetti | Messinian and pre-Messinian sediments from ODP Leg 107 Sites 652 and 654 in the Tyrrhenian Sea: sedimentologic and petrographic study and possible comparisons with Italian sequences[END_REF][START_REF] Cita | Messinian Paleoenvironments: New Evidence from the Tyrrhenian Sea (ODP LEG 107)[END_REF]Glaçon et al., 1990;Hasegawa et al., 1990;

Chapter 2-Material and Methods ___________________________________________________________________________________________ 80 Müller, 1990;Müller et al., 1990;Pierre and Rouchy, 1990;[START_REF] Müller | Origin and age of the Mediterranean Messinian evaporites: implications from Sr isotopes[END_REF]Flores et al., 1992;Roveri et al., 2014b;[START_REF] Lugli | The deep record of the Messinian salinity crisis: evidence of a non-desiccated Mediterranean Sea[END_REF].

Figure 2. 33 Core log at site 654 together with seismic reflection profile at this same site. From [START_REF] Kastens | The geological evolution of the Tyrrhenian Sea: An introduction to the scientific results of ODP Leg 107[END_REF] Due to extension and subsidence of the area inferred to have occurred after the Messinian Salinity Crisis, the Upper Gypsum from site 654 was most probably deposited in an intermediate depth basin, possibly shallower than at present day [START_REF] Loreto | Extensional tectonics during the Tyrrhenian back-arc basin formation and a new morpho-tectonic map[END_REF][START_REF] Lymer | The Western Tyrrhenian Sea revisited: New evidence for a rifted basin during the Messinian Salinity Crisis[END_REF]Roveri et al., 2014b). The Messinian unit sampled in this study (unit 2, cores 36 to 28) start with a rhythmic alternation of 8 primary cumulate gypsum (see [START_REF] Lugli | The deep record of the Messinian salinity crisis: evidence of a non-desiccated Mediterranean Sea[END_REF] for details) and marls beds which contain, at some stratigraphic levels, abundant Ammonia Kastens et al., 1987)(Figure 2. 31). Below, unit 3, also Messinian in age consists in dark, laminated, organic-rich, barren pelites but were not sampled in this study. on the seismic data [START_REF] Camerlenghi | Uncovering the Mediterranean Salt Giant (MEDSALT)-Scientific Networking as Incubator of Cross-disciplinary Research in Earth Sciences[END_REF]Hsü et al., 1978d) that marks the Messinian-Plio of halite interbedded with anhydrite (Figure 2. 31). This basin was probably shallower than nowadays because of Pliocene-Quaternary subsidence (Maillard et al., 2011;[START_REF] Reiche | The impact of salt on the late Messinian to recent tectonostratigraphic evolution of the Cyprus subduction zone[END_REF].

Site 374

Site 376

In powder and reacted with a reductive solution at ∼190°C under a continuous flux of nitrogen (Figure 2. 36). In order to obtain the purest H2S gas possible before its transformation into Ag2S, the products of the above-mentioned reaction were then pushed by the nitrogen flux through (i) a condenser that allows for the acid vapors by-products of the reaction to condense on the wall of the condenser, then (ii) through a water trap that allow residual gases to be solubilized (HCl for example). Finally, the resulting gas bubbled into 10mL of a sulfide trapping solution.

The sets of the experiment is similar to the one used in Oduro et al., 2011 (see Figure 2. 36).

Two types of reductives solution and sulfide trapping solutions were used, depending on the sulfur compounds that was targeted for the extraction.

Extraction of chromium reducible sulfur (iron sulfides) and transformation into Ag2S

A chromium solution (II) (CrCl2) allows for the extraction of sulfide minerals and S

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86

N2 until the colorimetric turning of the solution from its initial dark green to light blue (Fossing and Jørgensen, 1989).

Cr(II) is produced by the reaction of Cr(III) with Zn as follow :

2Cr 3+ + Zn 0 + H + +Cl -= Zn 2+ + 2Cr 2+ + H + + Cl -
Iron sulfide in the sample react with the hydrochloric acid to form H2S following this reaction:

FeS + 2HCl = Fe 2+ + H2S + 2Cl - Because S 2 -can be oxidized in S 0 in the presence of ferric iron Fe(III) (Pruden and Bloomfield, 1968), the presence of Cr(II) in the solution inhibited the oxidation of sulfur through :

2Cr 2+ + S 0 + 2H + = 2Cr 3+ + H2S This permitted the extraction and conversion of reduced sulfur compounds to H2S without reacting sulfates or organic sulfur (Gröger et al., 2009). This solution, being particularly sensitive to oxidation, was prepared on a daily basis according to Fossing and Jørgensen (1989). For each distillation, 20mL of this chromium reducible solution was injected together with 5mL of HCl (6N).

The H2S produced through this reaction was then reacted with a 2.48% AgNO3 sulfide trapping solution to produce a solid precipitate of silver sulfide through the following reaction:

H2S + 2Ag + + 2NO3 -= Ag2S + 2HNO3
The yield of the reaction for each sample was difficult to quantify since the initial quantity of sulfur in S 0 or iron sulfide form was unknown. However, repeated reaction of large amounts of gypsum or marls samples used in this work with excess heptane overnight never yielded any S 0 . We thus concluded that the sulfur extracted with the chromium reducible solution was mainly coming from the sulfide bearing species, most likely the iron sulfide minerals that were also observed by microscopic techniques (see chapters 3 and 4 for pictures and more details on this matter).

Extraction of sulfate from calcium sulfate minerals and transformation into Ag2S

For the reduction of sulfate, we used a "Thode" solution, adapted from [START_REF] Thode | Sulphur Isotope Geochemistry[END_REF], composed of HI, H3PO2 and HCl in order to convert the sulfate into H2S.

One common approach found in the literature is to first precipitate the targeted sulfate into barite (BaSO4) by reacting it with a BaCl2 solution. First because it is easily prepared from natural samples containing dissolved sulfate (e.g [START_REF] Johnston | Placing an upper limit on cryptic marine sulphur cycling[END_REF] and second because it can help separate the sulfate ion from other components produced during previous chemical steps, that may react together with sulfate during the acid extraction (e.g [START_REF] Jovovic | Quantitative and specific recovery of natural organic and mineral sulfur for (multi-)isotope analysis[END_REF]. The reaction at play would be the following: However, a recent study has demonstrated that the yield of such a reaction never reaches 100%, rather at maximum of 80%, because barite has a very low solubility that inhibits the reaction efficiency and speed [START_REF] Geng | A simple and reliable method reducing sulfate to sulfide for multiple sulfur isotope analysis[END_REF]. We thus directly used crushed gypsum or anhydrite that are more soluble than barite for the sulfate extraction with the Thode solution using the apparatus described above (Figure 2. 36). Given the very different concentration of sulfur from sulfate and sulfur from sulfide in the calcium sulfate minerals studied in this work, we didn't perform a sequential extraction of the sulfur compounds. Therefore, the sulfur extracted with the Thode solution represent the bulk of both sulfate and sulfide in our gypsum and anhydrite samples. However, the concentrations of sulfur in gypsum or anhydrite as sulfate are 10 3 to 10 6 larger than sulfur as sulfide. We thus consider the contribution of sulfide in the isotopic composition of the sulfate extracted with the Thode solution as negligible. The Thode solution was prepared by stirring 420 mL of HCl 37%, 257mL of HI 57% and 125mL of H3PO2 50% under a flux of N2 in order to homogenize it and to purify it from residual sulfate (Forrest and Newman, 1977;[START_REF] Thode | Sulphur Isotope Geochemistry[END_REF]. Reacting the H2S produced using Thode solution can lead to the formation of phosphorous compounds in addition to the targeted silver sulfide if directly reacted with AgNO3. To avoid this possible source for a bad chemical yield, the trapping solution used here was sodium hydroxide (NaOH 0.1M). At the end of the reaction, 3mL of AgNO3 were added to this solution, allowing the formation of Ag2O and of the targeted Ag2S. Few drops of HNO3 (68%) were then added to the solution that was shook to dissolve Ag2O while Ag2S remained in the solid phase. Samples with chemical yields below 80% or above 110% were not considered for isotope analysis.

Chapter 2-Material and Methods ___________________________________________________________________________________________ 89

SF6 production and purification

The Ag2S produced through the various extractions methods was rinsed 3 times, each time with 10mL of MilliQ water and centrifugated before being dried in an oven at 80°C overnight.

After this step, the production and purification of sulfur hexafluoride (SF6), the gas used for sulfur isotope analysis, was performed at the Stable Isotope Laboratory in IPGP, in a stainlesssteel vacuum line, for which a schematic representation is given in The produced SF6 has then to be purified before it is transferred to the mass spectrometer for isotopic analyses (Figure 2. 37).

(i)

The first step consists in the purification of the sample from excess F2 from the fluoraton step. F2 does not condense at -190°C in opposition to other gases including SF6. The content of the Ni bomb is thus sent into a variable temperature trap cooled at -190°C, the gas not condensed is sent to the passivation section to react with a KBr powder at 100°C that form BrF5.

(ii)

The temperature is then raised to -150°C. This allows the liberation of SF6 gas but leaves the undesirable HF into its liquid form.

(iii)

The released SF6 is condensed again in another liquid N2 trap before being liberated again and pushed by a helium flux (at about 10mL/min) to the Gas chromatography (GC) at 80°C. This allows the separation of the SF6 from the possible other gases (including C3F6) still presents in the sample.

(iv) Finally, the SF6 is separated from He by another liquid N2 trap and then transferred to an immersion sleeve to quantify the yield of the fluoration reaction before its transfer to the mass spectrometer.

Chapter 2-Material and Methods

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91

All samples having fluoration yields below 80% were not considered as valid datapoints.

Isotopic measurements

All the analyses have been performed on a gaseous dual inlet mass spectrometer

ThermoFinnigan MAT 253 at the Stable Isotope laboratory in IPGP (Au Yang et al., 2016;[START_REF] Labidi | Determination of multiple sulfur isotopes in glasses: A reappraisal of the MORB delta S-34[END_REF].

Figure 2. 38 Essential components of a dual inlet gas isotope ratio mass spectrometer with dual inlet system. Modified from Ghosh and Brand (2003) In the gaseous source of the mass spectrometer, tungsten filaments emit electrons that ionize the SF6 gas injected from one of the two gas reservoirs by the capillary. Ionization produces SF 5+ ions that are separated in the electromagnetic sector of the mass spectrometer depending on their mass/charge ratios, with masses 127, 128, 129 and 131 corresponding respectively to 32 S, 33 S, 34 S and 36 S sulfur isotopes. Because fluorine only has one isotope, if the analyzed gas is pure, there is no possible isobaric interference during measurements. The abundance of each isotope in the considered sample is then determined in the detection

__________________________________________________________________________________________ _

sector by Faraday cups. The dual inlet injection allows to alternate each sample measurement with an in-house SF6 standard measurement. The in-house SF6 standard is of a known isotopic composition. Each data is the result of 10 cycles of analyses standardsample. The isotopic compositions in δ values were calibrated with the in-house standard using IAEA-S1 international standard that has a δ 34 S value of -0.3‰ VS-CDT. Isotopic compositions are then reported in V-CDT notation (Vienna Canyon Diablo Troilite).

To ensure the quality of the measurement , each measurement session begins with a zero enrichment, where the internal reference gas is injected in the two reservoirs of the mass spectrometer. The reference gas being measured against itself, the result of such measurement should yield 0.000‰. The zero e nrichment (n=22) yielded δ 33 =-0.005±0.007‰, δ 34 =-0.004±0.006‰, δ 36 =-0.005±0.071‰ respectively during the period where IODP and DSDP samples have been measured (n=24) δ 33 =0.000±0.005‰, δ 34 =0.001±0.01‰, and δ 36 =0.007 ±0.059‰ during the period where measurements of Vena del Gesso samples have been carried out. This demonstrates that no significant residual pollution was susceptible to disturb our measurements.

On the other hand, long term reproducibility of international standards IAEA-S1, IAEA-S2 and IAEA-S3 allowed to ensure of the good functioning of the SF6 purification line. On average, the measurements of these standards gave results very close to the reported reference for δ 34 S [START_REF] Brand | Assessment of international reference materials for isotope-ratio analysis (IUPAC Technical Report)[END_REF] [START_REF] Pierre | The isotopic record of gypsum diagenesis in diluted solutions: Observations in natural salinas and experiments[END_REF]. The δ 18 O values were calibrated using internal laboratory standard Edm-Tyr (δ 18 O =+8.6‰ V-SMOW) that was calibrated against the IAEA international reference NBS 127 (δ 18 O = +8.5 ± 0.2‰ V-SMOW, Boschetti and Iacumin, 2005). The analytical precision is 0.2‰.

Calibration of the isotopic data to international standards

δ 34 S and δ 18 O from this study have been set to the VSCDT and VSMOW scale respectively using recent calibration of NBS-127 reference material [START_REF] Brand | Assessment of international reference materials for isotope-ratio analysis (IUPAC Technical Report)[END_REF].

Chapter 2-Material and Methods

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94

To compare these published data with ours, we corrected the reported data that were calibrated with obsolete values of NBS-127 by using the difference between the reported and newly calibrated δ 34 S and δ 18 O values of the standard following the method reported in [START_REF] Johnston | Placing an upper limit on cryptic marine sulphur cycling[END_REF]. In addition, because there is a small and almost temperatureindependent isotopic fractionation between the dissolved SO4 2-ion and gypsum, we calculated the isotopic composition of the sulfate ion in equilibrium with gypsum by applying known fractionation factors (at earth surface temperature α 34 Gyp-SO4=1.00165 [START_REF] Raab | Sulfur isotopic variations during seawater evaporation with fractional crystallization[END_REF]Thode and Monster, 1965;[START_REF] Van Driessche | Unraveling the Sulfate Sources of (Giant) Gypsum Crystals Using Gypsum Isotope Fractionation Factors[END_REF] ; 33 αGyp-SO4=1.00084 and α 18 Gyp- SO4=1.0035 (Lloyd, 1968;[START_REF] Van Driessche | Unraveling the Sulfate Sources of (Giant) Gypsum Crystals Using Gypsum Isotope Fractionation Factors[END_REF]).

Error calculation for lambda fractionation factors

Fractionation factors εnet and λnet between sulfide and sulfate have been calculated (see section 1.9 for mathematical expressions). These calculations were carried out in order to estimate the net isotopic effect of the processes at play during the formation of our samples.

There is a dependence between the σλnet error and the magnitude of the fractionation, εnet (Johnston et al., 2007). To illustrate this dependence we calculated the σλnet incertitude using the expression from Bevington and Robinson (2003) following Johnston et al. (2007) : The uncertainties calculated as described above account for the dependence of σλ on the magnitude of fractionation ε as shown in Figure 2. 39, and were used to calculate error bars for the λnet from our samples. In most natural environments, more than one process and/or pole of sulfur are involved and it is not straightforward to decipher their relative influences in complex systems. In Chapter 3, we use multiple sulfur isotope data to determine the processes and the pool of sulfur used during the formation of the Vena del Gesso samples. To this end, we designed a numerical model to quantify the relative fluxes of sulfur involved. The following section is a description of this numerical model, that is mainly issued from the supplementary materials of the paper from which Chapter 3 comes from (Guibourdenche et al., 2022).

𝜎 "#$% = ' 𝜎

Sulfur fluxes and biogeochemical processes

The steady-state, 3-box model (Figure 2. 40) is based on a chemical and isotopic mass balance of sulfur in the SO4

___________________________________________________________________________________________

98

In summary, the sulfur isotope composition of gypsum and pyrite accumulating in a marginal basin can vary as a function of (i) mixing between the ocean-and river-derived sulfate sources, (ii) the magnitude of gypsum and pyrite burial fluxes and (iii) the extent and mode of intrabasinal biogeochemical sulfur cycling and the associated isotope fractionation. Within the marginal basin, sulfur is abiotically and microbiologically partitioned between the three sulfur pools SO4 2-, S° and H2S via the following processes:

(iii.2) biologically, via phototrophic sulfide oxidation (green or purple sulfur bacteria):

4CO2 + 2H2S + 4H2O → 4CH2O + 2SO4 2-+ 4H + (Findlay et al., 2017) (6) Or 2HS -+ CO2 + H + → 2S 0 + CH2O + H2O (Findlay et al., 2017) (6.1) and Chapter 2-Material and Methods ___________________________________________________________________________________________ 100 3CO2 + 2S 0 + 5H2O → 3CH2O + 2SO4 2-+ 4H + (Findlay et al., 2017) (6.2) (iii.3) biologically, via chemolithotrophic sulfide oxidation (colorless sulfur bacteria) using O2 or NO3 -as electron acceptors:

HS -+ NO3 -+ H + + H2O → SO4 2-+ NH4 + (7)
Or 4HS -+ NO3 -+ 6H + → 4S In our model, the oxidation of H2S takes place either via processes 3,4,5,6 or 7 (with SO4 2-as an end product) or via processes 3.1, 6.1 or 7.1 (with S 0 or SO3 2-as an end product). While the rate of all of these four processes is denoted Fox, each one of the four processes implies a distinct sulfur isotope fractionation, as defined in Table 2. 1.

If not reoxidated into sulfate, the produced intermediate sulfur compound (S° or SO3 2-) can be involved in:

IV. disproportionation of S 0 or SO3 2-producing sulfate and sulfide, with rates Fdo and Fdh, respectively: and Cypionka, 1987;Thamdrup et al., 1993)

4S 0 + 4H2O → 3H2S + SO4 2-+ 2H + (Bak
Or 4SO3

Finally, aqueous sulfide produced by microbial reactions can be extracted from the marginal basin by:

V. solid iron sulfide formation and burial (Fcrs): through aqueous sulfide reaction with iron, it is, in our case, difficult to assess the exact nature of the CRS preserved in our samples that we extracted chemically. Moreover, it is most likely that we are in presence of different types of solid sulfide compounds possibly formed by different processes. Therefore, and given that the fractionation factors remain very small and are most likely to have a negligible effect on the isotopic signature of the CRS in our samples (Eldridge et al., 2021;Fry et al., 1986), we consider the fractionation during solid sulfide burial as negligible (see Table 2. 1).

HS -+ Fe 2+ → FeS + H + (

Inorganic oxidation :

Because there are no known fractionation factors for abiotic sulfide oxidation by iron and manganese oxides, we considered, like Pellerin et al. (2015), that these processes fractionate sulfur isotopes similar to abiotic oxidation of sulfide by molecular oxygen (Eldridge and Farquhar (2018)).

Variable fractionation factor for microbial sulfate reduction

Multiple pure culture and environmental studies have shown that the magnitude of sulfur isotope fractionation during MSR depends strongly on the availability of electron donor or acceptor, as they affect the cellular rate of microbial sulfate reduction (csSRR, in mol of sulfide produced per day per unit biomass) [START_REF] Bradley | Patterns of sulfur isotope fractionation during microbial sulfate reduction[END_REF]Habicht and Canfield, 2001;Leavitt et al., 2013;Sim et al., 2011a;Sim et al., 2011b). These authors found an inverse correlation between csSRR and the magnitude of isotopic fractionation. Leavitt et al. 2013, applied a nonlinear regression model (equations ( 12) and ( 13 where 34 ƐMeasured and 33 λMeasured represent the extent of fractionation measured for each individual SO4 2--H2S couple during the pure culture experiment, whereas 34 ƐMin , 33 λMin , 34 ƐMax and 33 λMax are the minimum and maximum empirical fractionations limits and 34 kmsr and 33 kmsr are pseudo first order rate constants obtained by fitting equations ( 12) and ( 13) to the experimental data.

To express this first order dependence of csSRR on sulfur isotope fractionation in our model, we rearranged equations ( 12) and ( 13) to obtain expressions for 34 ƐModelled and 33 λModelled as function of csSRR:

{| 𝜀 }~$ÄÄ$ = 𝑒 9ÇÉ+ÑÑ× {| x yqs (𝜀 }ÖÜ -𝜀 }á# ) + 𝜀 }á# (àâ)
{{ 𝜆 }~$ÄÄ$ = 𝑒 9ÇÉ+ÑÑ× {{ x yqs (𝜆 }ÖÜ -𝜆 }á# ) + 𝜆 }á# (àä) Leavitt et al., 2013, however, explored only part of the potential fractionations possibly yielded by MSR communities because they only used lactate as an electron donor and set the sulfate concentration constant. On the other hand, Sim et al. (2011) carried out pure culture experiments with a variety of electron donors and sulfate concentrations, observing a wider range of isotope fractionations that approach the thermodynamic equilibrium between SO4 2- and H2S equal to 70‰ (Brunner and Bernasconi, 2005;Farquhar et al., 2003;Tudge and Thode, 1950).

In our model, in order to cover the full range of isotopic fractionations that are possibly expressed by MSR activity, we expressed the cellular MSR rate of the Sim et al., 2011aSim et al., , 2011b experiments on a normalized scale, where csSRR= 0 corresponds to no sulfate reduction and

Chapter 2-Material and Methods ___________________________________________________________________________________________ 105 csSRR=100 corresponds to the rate of sulfate reduction where the minimum isotope fractionation is expressed. We then refitted the data of Sim et al., 2011aSim et al., , 2011b to equations ( 14) and ( 15) (Figure 2. 41). This allowed us to obtain our own 34 ƐMin (-69) , 33 λMin (0.511) , 34 ƐMax (-5) , 33 λMax (0.515), 34 kmsr (0.0514) and 33 kmsr (0.1409) values.

66 𝜆 àÉâ (ε àÉâ ) = (𝑒 9ÇÉ+ÑÑ(ε yqs )× ** x yqs (𝜆 àÖÜ -𝜆 àá# ) + 𝜆 àá# (àã)

{{ 𝛼 àÉâ = {| 𝛼 àÉâ (ε àÉâ )

{{ " yqs (ε)

(àå)

In the model λ and α were calculated as function of csSRR that was modulated by the entire range of possible values for εMSR (-ä to -åç‰).

Model equations

The model is based on two mass balances for each of the three boxes (Figure 2. 40): one chemical mass balance for total sulfur and one isotopic mass balance for the rare sulfur isotopes (the isotope mass balance equations are equivalent for 34 S and 33 S).

Where fluxes (Fx) are defined as in and Rx are the isotope ratios ( x S/ 32 S, where x = 34 or 33) of the same fluxes.

To obtain output fluxes of gypsum-and pyrite-S expressed as a percentage of the sulfur fluxes added to the marginal basin from external sources, we define:

𝐹 á# + 𝐹 âáå = 100 (20)
Furthermore, we define parameters (fx, varying from 0 to 1) that express the relative magnitude of the different sulfur fluxes. Thus, fex represents the fraction of the incoming sulfur flux that is not involved in microbial reactions or gypsum precipitation:

𝑓 $Ü = ì îrM (ì uK Pì suï ) (21)
friv designates the fraction of the external sulfur inputs that is brought to the marginal basin by rivers:

𝑓 âáå = ì suï (ì uK Pì suï ) ( 22
)
and fmsr is the fraction of the net incoming sulfur flux being removed from the marginal basin through microbial sulfate reduction rather than by gypsum accumulation:

𝑓 àÉâ = ì yqs (ì uK Pì suï 9ì îrM ) (23)
For the H2S box, the two mass balances are:

For the H2S box, we define fox as the fraction of the incoming sulfur fluxes that leaves the H2S box through H2S -oxidation (reactions 3, 4, 5 or 6) instead of being precipitated as pyrite:

𝑓𝑜𝑥 = ì îv (ì yqs Pì tô ) (26)
For the S° box, the two mass balances are:

𝐹 ~Ü = 𝐹 ñ + 𝐹 ~ (27)
And 𝐹 ~Ü × 𝑅 ~Ü = 𝐹 ñ × 𝑅 ñ + 𝐹 ~× 𝑅 ~ [START_REF] Aloisi | The geochemical riddle of ''low-salinity gypsum" deposits[END_REF] For the S° box, we define fdh as the fraction of sulfur added to the S° box by oxidation of H2S that is re-reduced to H2S through disproportionation:

𝑓𝑑ℎ = ì tô ì îv (29)
Finally, we define isotope fractionation factors for the considered processes that express a Sisotope fractionation:

𝛼 çéè = Ñ úùû Ñ yü (30)
is the fractionation factor associated to the precipitation of gypsum,

𝛼 àÉâ = Ñ yqs Ñ yü (31)
Chapter 2-Material and Methods ___________________________________________________________________________________________ 108 is the fractionation factor associated to microbial sulfate reduction

𝛼 èé = Ñ ûù Ñ ,-. (32)
is the fractionation factor associated to pyrite formation

𝛼 ~Ü = Ñ îv Ñ ,-. (33)
is the fractionation factor associated to H2S oxidation

𝛼 ñ = Ñ tô Ñ .° (34)
is the fractionation factor associated to reduction of S° to H2S through S° disproportionation

𝛼 ~= Ñ tî Ñ .° (35)
is the fractionation factor associated to oxidation of S° to SO42-through S° disproportionation

The set of 18 equations (18 to 35) and 18 unknowns (Fin,Fox,Friv,Fout,Fgyp,Fcrs,Fmsr,Fdh,Fdo,RSO4,Rgyp,Rpy,RH2S,Rox,Rmsr,Rdo,Rdh,Rs°) was solved analytically with the Wolfram Mathematica 11 software, providing mathematical expressions for all sulfur fluxes and their isotope ratios (equations 40 to 57, below).

To facilitate reading of equations ( 40) to (57), we define:

𝐴 = 𝛼 çéè (𝑓 $Ü -1)(1 + 𝑓 àÉâ (𝑓 ~Ü -1) -𝑓 ñ 𝑓 ~Ü ) (36) 𝐵 = -𝑓 $Ü + 𝛼 àÉâ 𝑓 àÉâ (𝑓 $Ü -1)(𝑓 ñ 𝑓 ~Ü -1) (37) 𝐶 = 𝛼 èé (𝑓 ~Ü -1) (38) 𝐷 = (𝑓 ñ 𝑓 ~Ü -1)(𝑅 ~Ç(𝑓 âáå -1) -𝛼 âáå 𝑅 âáå ) (39)
Following these new notations, solutions of equations ( 18) to (35) are given below

𝐹 á# = -100(𝑓 âáå -1) (40)
𝐹 âáå = 100𝑓 âáå (41)

𝐹 ~ê% = 100 × 𝑓 $Ü ( 42
)
Chapter 2-Material and Methods

___________________________________________________________________________________________

109

𝐹 àÉâ = -100𝑓 àÉâ (𝑓 $Ü -1) (43) 𝐹 ~Ü = - ì yqs • îv • tô • îv 9O (44)
𝐹 ñ = 𝐹 ~Ü𝑓 ñ (45)

𝐹 ~= -𝐹 ~Ü(𝑓 ñ -1) (46)
𝐹 çéè = - ì îv (• yqs P• tô • îv 9• yqs • îv 9O) • yqs • îv (47) 𝐹 èé = ì yqs (• îv 9O) (• tô • îv 9O) (48) 𝑅 +:| = ¶[9® tô • tô ©P® tî (• tô 9O)(©9® îv • îv )] [® tô • tô ©(´9¨)] P ® tî (• tô 9O)[(¨©P® îv • îv • Lv (• tô • îv 9O)9´(©9® îv • îv )]
(âé)

𝑅 àÉâ = 𝑅 +:; 𝛼 àÉâ (50)
𝑅 çéè = 𝑅 +:; 𝛼 çéè (51)

𝑅 8(+ = ¶® yqs (• tô • îv 9O)[® tî (• tô 9O)9® tô • tô ] [® tô • tô ©(´9¨)] P ® tî (• tô 9O)[(¨©P® îv • îv • Lv (• tô • îv 9O)9´(©9® îv • îv )] (52)
𝑅 èé = 𝑅

𝑅 +°= - Ñ ,-. ® îv [® tî (• tô 9O)9® tô • tô ] (57)
To this set of equations were added equations ( 16) and ( 17) that allowed the MSR fractionation to vary as a function of csSRR.

Introduction

At the end of the Miocene, starting from about 7.2 Ma, the progressive reduction of water exchanges between the Atlantic Ocean and the Mediterranean Sea dramatically affected Mediterranean environments, increasing their sensitivity to orbitally-driven climate perturbations (Bulian et al., 2021;Hsü et al., 1973;[START_REF] Kouwenhoven | Paleoenvironmental evolution of the eastern Mediterranean during the Messinian: constraints from integrated microfossil data of the Pissouri Basin (Cyprus)[END_REF]. Hydrological restriction to the Mediterranean led to an increase of salinity, a decrease in water column oxygenation and changes in biotic assemblages (Bulian et al., 2021;Cita and Ryan, 1978;[START_REF] Kouwenhoven | Paleoenvironmental evolution of the eastern Mediterranean during the Messinian: constraints from integrated microfossil data of the Pissouri Basin (Cyprus)[END_REF]. Thereafter, the sensitivity of the Mediterranean to precessional cycles became significantly more pronounced, and in turn greatly impacted the biogeochemistry of the basins (Natalicchio et al., 2019;[START_REF] Sabino | The response of water column and sedimentary environments to the advent of the Messinian salinity crisis: insights from an onshore deep-water section (Govone, NW Italy)[END_REF]. This phenomenon is recorded as cyclical deposition of sapropel/diatomite-rich marls (Krijgsman et al., 1999) and marl/carbonate alternations [START_REF] Mancini | Calcareous nannofossil size and abundance response to the Messinian salinity crisis onset and paleoenvironmental dynamics[END_REF] corresponding to periods of summer insolation maxima and minima, respectively. Lithological cyclicity is accompanied by variations in the size and abundance of fossils of marine fauna [START_REF] Mancini | Calcareous nannofossil size and abundance response to the Messinian salinity crisis onset and paleoenvironmental dynamics[END_REF], as well as oscillations in the composition and isotopic signature of organic lipid biomarkers (Natalicchio et al., 2019;[START_REF] Sabino | Climatic and hydrologic variability in the northern Mediterranean across the onset of the Messinian salinity crisis[END_REF]. Together, this evidence suggests the existence of a precessional-scale alternation between a highly productive/stratified water column and a vertically mixed water column.

The progressive intensification of hydrological restrictions between the Atlantic and the Mediterranean culminated at 5.97 Myrs with the onset of evaporite precipitation that marked the beginning of the Messinian Salinity Crisis (MSC). During the first phase of the MSC, up to 220 m of cyclical alternations of gypsum and organic-rich deposits (characteristic of arid and humid periods, respectively) forming the Primary Lower Gypsum (PLG) unit (Figure 3. 1)

Chapter 3-Cryptic sulfur cycling during the formation of giant gypsum deposits _____________________________________________________________________________________________________ 114 accumulated in the shallow part of marginal to intermediate depth basins of the Mediterranean (Lugli et al., 2010;Raad et al., 2021;Roveri et al., 2014). During this time, marginal and intermediate depth basins still received significant marine water inputs, as indicated by the 87 Sr/ 86 Sr and δ 34 S of gypsum and by the occurrence of marine biota (Lugli et al., 2007;Lugli et al., 2010 and references therein;Pellegrino et al., 2021;Reghizzi et al., 2018). [START_REF] Laskar | La2010: a new orbital solution for the long-term motion of the Earth[END_REF]. Note the nearly invariant δ 34 SSO4 isotopic variations compared to δ 34 SIS. B.

Panoramic view of the 220m thick VdG succession constituted by an alternation of thick gypsum beds and marl intervals (PLG unit). C. Distribution of Messinian evaporites in the Mediterranean and the location of the VdG basin (pink star) (see Chapter 2 for references).

During humid periods of precession cycles, non-evaporitic interbeds deposited in stratified basins where N2-fixing bacteria significantly contributed to primary productivity (Isaji et al., 2019) and anoxygenic phototrophic bacteria populated the upper water column (Sinninghe Damsté et al., 1995). Meanwhile, microbial mats dominated by colorless sulfide oxidizing Chapter 3-Cryptic sulfur cycling during the formation of giant gypsum deposits _____________________________________________________________________________________________________ 115 bacteria (CSB) thrived at the seafloor (Dela Pierre et al., 2014). Anoxygenic phototrophic bacteria and CSB mats develop in modern environments in the presence of H2S fluxes produced by microbial sulfate reduction in the euphotic zone and the benthic boundary, respectively. This H2S is oxidized back to elemental sulfur or sulfate [START_REF] Brüchert | Biogeochemical and Physical Control on Shelf Anoxia and Water Column Hydrogen Sulphide in the Benguela Coastal Upwelling System off Namibia[END_REF]Callbeck et al., 2021;[START_REF] Jørgensen | Sulfur biogeochemical cycle of marine sediments[END_REF]. The presence of sulfide oxidizing metabolisms in both the water column and at the sediment-water interface during marls/organic rich shale deposition strongly suggest the prevalence of euxinic conditions.

Precession maxima (arid periods) were characterized by rapid vertical accretion of bottom grown gypsum crystals (Reghizzi et al., 2018) -which allowed for the relatively good preservation of organic remains (Dela Pierre et al., 2015;Schopf et al., 2012) (Figure 3. 2).

These bottom-grown gypsum crystals make up the bulk of the PLG unit, and contain abundant planktic marine diatoms, implying enhanced primary productivity in the upper water column during gypsum deposition (Dela Pierre et al., 2015;Pellegrino et al., 2021;Schopf et al., 2012).

The molecular fossil assemblages entrapped within Messinian gypsum crystals and their compound specific carbon isotopes are indicative of a diverse microbial community, including marine Thaumarchaeota communities living in the upper water column, and sulfate-reducing bacteria and CSB-forming microbial mats at the sea floor (Natalicchio et al., 2022). The latter benthic microbial mats are fossilized within the bottom grown gypsum in the form of denselypacked networks of filamentous structures ("spaghetti-like" structures) together with pyrite and dolomite, implying that microbial sulfate reduction activity occurred during gypsum formation (Dela Pierre et al., 2015;Natalicchio et al., 2022;[START_REF] Panieri | Microbial communities in Messinian evaporite deposits of the Vena del Gesso (northern Apennines, Italy)[END_REF]Schopf et al., 2012) (Figure 3. 2D-I).

Chapter 3-Cryptic sulfur cycling during the formation of giant gypsum deposits _____________________________________________________________________________________________________ 116 Together, these petrographic and geochemical observations point towards a higly dynamic sulfur cycle during PLG formation, with sulfur cycled amongst microbial sulfate reducers and sulfide-oxidizing bacteria, as well as precipitated in its oxidized form as gypsum (CaSO4.2H2O).

Marine carbon and sulfur cycles are tightly linked, with MSR driving the remineralization of between 12 and 65% of the organic carbon flux to the seafloor [START_REF] Bowles | Global rates of marine sulfate reduction and implications for sub-sea-floor metabolic activities[END_REF][START_REF] Bradley | Widespread energy limitation to life in global subseafloor sediments[END_REF]Canfield et al., 2005). Constraining sulfur cycling during the formation of giant sulfate deposits is therefore critical to understand their significance for the carbon cycle. The accumulation of large gypsum deposits -comparable to the Mediterranean PLG-are known to greatly affect the long term carbon cycle and atmospheric oxygen levels by the removal of substantial amount of Ca 2+ and SO4 2-ions from the ocean, (e.g. Halevy et al., 2012;Shields and Mills, 2020). However, the role of associated MSR and sulfide oxidation on carbon and sulfur cycle of basins where gypsum formed has been overlooked.

The complexity of sulfur cycling dynamics in marginal basins during PLG formation (which were likely to have been widespread across the Mediterranean (Natalicchio et al., 2022;Pellegrino et al., 2021)), raises fundamental questions about the dynamics of these basins, including for instance: (i) What was the relative importance of microbially driven reduction of sulfate and reoxidation of sulfide, compared to the burial of sulfate through gypsum precipitation ? (ii) How did sulfur cycling impact carbon cycling in Messinian Mediterranean marginal basins? (iii) How does this ancient system compare with modern marine environments ?

In this study, we investigated the most complete PLG succession of the Mediterranean area in the Vena del Gesso basin (VdG) (Northern Apennines, Italy) (Figure 3. 1B). MSR and reoxidative sulfur cycling result in specific multiple sulfur isotopes signatures, providing a tool Chapter 3-Cryptic sulfur cycling during the formation of giant gypsum deposits _____________________________________________________________________________________________________ 117 to disentangle the relative influence of concurrent sulfur redox processes [START_REF] Johnston | Sulphur isotopes and the search for life: strategies for identifying sulphur metabolisms in the rock record and beyond[END_REF]Ono et al., 2006). Therefore, to constrain the sulfur biogeochemical cycle during PLG deposition, we carried out measurements of multiple sulfur isotope ratios ( 34 S/ 32 S and 33 S/ 32 S) on gypsum sulfate and reduced sulfur minerals which co-precipitated with gypsum or that we extracted from marls, coupled to the measurement of oxygen isotopes of gypsum sulfate ( 18 O/ 16 O). This approach enabled us to estimate the relative fluxes of sulfide and sulfate in the basin and in turn to estimate the gross flux of organic carbon involved in MSR activity. In modern settings, this is typically achieved with incubation experiments and/or analyses of pore water geochemical profiles -tools that are not available when studying the rock record.

Geological setting

The PLG was deposited during the first stage of the MSC in shallow to intermediate marginal basins of the Mediterranean, including southern Spain, Italy, Hellenic arc and Cyprus arc (Krijgsman et al., 1999;Lugli et al., 2010 and references therein) (Figure 3. 1C). This unit was deposited in approximately 380 kyrs (Roveri et al., 2014) and is formed by lithological cycles consisting of organic-rich (marls, shale and carbonate)/gypsum couplets (Figure 3. 1A). The deposition of these interbeds was controlled by precession-driven climate oscillations (Lugli et al., 2010;Roveri et al., 2014). Tuning of the lithological cycles to the astronomical reference curve (Figure 3. 1A) allows the precise dating of each cycle and bed by bed correlation of the PLG unit at the Mediterranean scale (see the pink dots in Figure 3. 1C) (Krijgsman et al., 1999;Roveri et al., 2014). Here we studied the VdG in the Northern Apennines (Italy) that exposes the most complete succession of PLG deposits in the whole Mediterranean area (Figure 3. 1B). Each PLG cycle of the VdG basin starts with a thin (dm to Chapter 3-Cryptic sulfur cycling during the formation of giant gypsum deposits _____________________________________________________________________________________________________ 118 m thick) organic matter-rich (total organic carbon from 0.2 to 3.1%; Lugli et al., 2007) marly bed (Figure 3. 2A).

The thin marly beds are followed in each PLG cycle by a 4 to 30 m thick bed of primary gypsum (Figure 3. 2A). The gypsum beds of the first three cycles (PLG1 to PLG3) are mostly composed of bottom-grown, vertically oriented, twinned selenite crystals (up to 2.5 m high) of the massive selenite lithofacies (Lugli et al., 2010). Starting from the gypsum beds of the 4th PLG cycle, the massive selenite lithofacies is followed upwards by the banded selenite lithofacies composed of less than 10 cm-thick crusts of bottom-grown twinned gypsum crystals separated by thin carbonate laminae possibly reflecting cyclical fluctuation of the gypsum saturation interface (Lugli et al., 2010) (Figure 3.

2B).

From the 6 th PLG cycle to the top of the section, massive and banded selenite are followed by the branching selenite lithofacies consisting of cm-sized clear bottom-grown selenite crystals growing inclined or horizontally and grouped together in nodules or lenses separated by thin fine grained carbonate or gypsum laminae (Lugli et al., 2010). The absence of karst features in any of these lithofacies suggests that PLG gypsum are fully subaqueous deposits formed in a basin permanently covered by waters supersaturated with respect to gypsum. The sulfur and oxygen stable isotope compositions of pore water dissolved sulfate are frequently used to investigate the biogeochemical sulfur cycling in modern marine sediments (Aller et al., 2010;Blonder et al., 2017;[START_REF] Gilhooly | A comprehensive sulfur and oxygen isotope study of sulfur cycling in a shallow, hyper-euxinic meromictic lake[END_REF]. These stable isotopic 3. 3) (Brunner et al., 2005).

Isotopes to decipher sulfur cycling

However, the effect of MSR on sulfate and hydrogen sulfide concentration profiles and isotopic compositions can be obscured by re-oxidative processes; this happens because 34 Sdepleted sulfur metabolized by MSR is re-injected back into the 34 S-enriched sulfate pool with little or no fractionation (Figure 3. 3). This results in a smaller net enrichment of dissolved sulfate δ 34 S due to microbial sulfate reduction, and a lower net production of sulfide/consumption of sulfate. Importantly, if all the hydrogen sulfide produced by MSR is entirely re-oxidized back to sulfate (Figure 3. 3) (Aller et al., 2010;Blonder et al., 2017;[START_REF] Brüchert | Biogeochemical and Physical Control on Shelf Anoxia and Water Column Hydrogen Sulphide in the Benguela Coastal Upwelling System off Namibia[END_REF]Callbeck et al., 2021), the sulfur cycle is "cryptic" -since there is no net isotopic effect on the δ 34 S isotopic composition. Furthermore, the shift in δ 18 O sulfate isotopic values is often so small (<4‰) (Aller et al., 2010;Blonder et al., 2017) that it can be difficult to recognize the effect of cryptic sulfur cycling from natural variability of oceanic sulfate isotopic composition using this isotope system (Figure 3. 3) (Markovic et al., 2016). [START_REF] Evans | Coupled measurements of delta O-18 and delta D of hydration water and salinity of fluid inclusions in gypsum from the Messinian Yesares Member, Sorbas Basin (SE Spain)[END_REF]García-Veigas et al., 2018), Bajo Segura (García-Veigas et al., 2018), Murcia (Pierre, 1982) and Mallorca (García-Veigas et al., 2018), the Italian marginal Piedmont and Caltanissetta basins (El Kilany, 2018;García-Veigas et al., 2018) and the Polemi and Kalavasos/Psematismenos basins of Cyprus (Pierre, 1982)) (see section 2.3.3 for calibration of δ 34 SSO4 and δ 18 OSO4 isotopic data from this study and from the literature). The range of isotope compositions of sulfate-bearing minerals (Boschetti et al., 2011 and references therein;[START_REF] Claypool | The age curves of sulfur and oxygen isotopes in marine sulfate and their mutual interpretation[END_REF]Crockford et al., 2019;Ortí et al., 2010;Sakai, 1971;Utrilla et al., 1992) and of sedimentary sulfides (Balci et al., 2007;Burke et al., 2018) are represented as grey and light brown boxes. The possible range of isotopic composition of riverine inputs is bracketed by the isotopic composition of sulfate-bearing minerals and sedimentary sulfides representing the riverine endmember. Note that the contribution of riverine SO4 to gypsum SO4 isotopic composition seems negligible. The possible evolution of the isotopic composition of residual sulfate during MSR is indicated by the orange shaded area [START_REF] Fotherby | Modelling the Effects of Non-Steady State Transport Dynamics on the Sulfur and Oxygen Isotope Composition of Sulfate in Sedimentary Pore Fluids[END_REF]. The isotope composition of sulfate dissolved in top few centimeters of modern marine environments characterized by near complete reoxidation of H2S is represented by the blue box (Aller et al., 2010).

Chapter 3-Cryptic sulfur cycling during the formation of giant gypsum deposits _____________________________________________________________________________________________________ 122

Multiple sulfur isotopes

Multiple sulfur isotopes (including 32 S, 33 S, 34 S) have the potential to distinguish biogeochemically distinct hydrogen sulfide re-oxidation pathways occurring concomitantly with MSR -the signal of which might otherwise be obscured using traditional isotope systems [START_REF] Johnston | Sulphur isotopes and the search for life: strategies for identifying sulphur metabolisms in the rock record and beyond[END_REF]Ono et al., 2006;[START_REF] Zhang | Redox chemistry changes in the Panthalassic Ocean linked to the end-Permian mass extinction and delayed Early Triassic biotic recovery[END_REF].

The partitioning of 34 S ( 34 S/ 32 S) relative to 33 S ( 33 S/ 32 S) is predictable under thermodynamic equilibrium. However, intracellular biochemical branching processes will shift the minor 33 S sulfur isotope partitioning ( 33 S/ 32 S) thus resulting in a deviation from the expected thermodynamic equilibrium relationship (Ono et al., 2006). This deviation is expressed as Δ 33 S where:

∆ 77 𝑆 9 = :& << ' ()*+,-). << ' =>?@ -13 × 1000A -B& <C ' ()*+,-). <C ' =>?@ -13 D )E × 1000F (2)
The measurement of δ 34 S-Δ 33 S thus provides additional information as deviation from thermodynamic equilibrium can be linked to biological fractionation processes and mixing proportions of end-member sulfur species involved in sulfur cycling (Figure 3. 4) (Ono et al., 2006).

The deviation from thermodynamic equilibrium is also expressed by the 33 λ notation that defines the relationship between isotopic fractionation for 33 S and 34 S isotopes among two sulfur species:

77 𝜆 H = IJ( << K L ) IJ( <C K L ) (3)
with λ for thermodynamic equilibrium 33 λeq = 0.515, and x αj = x/32 Rproduct/ x/32 Rreactant, and x = 33 or 34.

Combining the 33 λ and the classical 34 Ɛ [= ( 34 αj -1) × 1000] notations is particularly useful as λ and ε fractionation factors are dependent solely on the biogeochemical processes in question Chapter 3-Cryptic sulfur cycling during the formation of giant gypsum deposits _____________________________________________________________________________________________________ 123 (e.g. MSR, biotic or abiotic sulfide oxidation). In addition, the combination of 33 λ and 34 Ɛ is not sensitive to the absolute isotopic value of the reacting sulfur species (Figure 3. 5) or the almost invariant temperatures relevant for sedimentary environments [START_REF] Eldridge | Theoretical estimates of equilibrium sulfur isotope effects in aqueous sulfur systems: highlighting the role of isomers in the sulfite and sulfoxylate systems[END_REF].

Material and Methods

A total of 88 samples that are representative of each lithofacies of all cycles from the VdG succession were analyzed (Figure 3. 1A). Gypsum and marls from PLG cycle 1-2 were sampled in the Monticino quarry (44°13ʹ29ʺN, 11°45ʹ43ʺE) whereas cycles 3 to 16 were sampled in the Monte Tondo quarry (44°15ʹ04ʺN, 11°40ʹ13ʺE). Care was taken to avoid sampling the weathered surface and visible fractures in order to limit contamination.

Petrographic analyses

11 oriented samples were taken from the various gypsum lithofacies for petrographic observations. Fifteen petrographic thin sections and 10 thin slices (~1 mm thick) were produced from the most representative samples by cleaving the selenite crystals along the main (010) cleavage planes with a razor blade. These were studied using transmitted, reflected, and UV-visible light microscopy, using a Leika DM 2700 P microscope with excitation filters 340-380 Nm at the Department of Earth Sciences of the University of Torino.

Oxygen and multiple sulfur isotope analyses

Sulfate from gypsum and iron sulfide from gypsum and marls were first reduced and transformed into Ag2S by wet chemistry. The multiple sulfur isotope composition was then measured using a MAT-253 mass spectrometer. Oxygen isotopic composition of sulfate from Chapter 3-Cryptic sulfur cycling during the formation of giant gypsum deposits _____________________________________________________________________________________________________ 124 gypsum were measured on samples precipitated as BaSO4 (the full methods with references are detailed in section 2.3).

Geochemical modeling

Description of the model

We interpret the multiple S isotope data using a steady-state, 3-box model based on a chemical and isotopic mass balance of sulfur in the SO4 2-, S° and H2S forms. In the model (see Figure 3. 6) sulfate is provided by oceanic (Fin) and riverine inputs (Friv) whereas it is removed from the basin through outflow into the open ocean (Fout) or gypsum precipitation (Fgyp). In addition, the sulfate can be reduced to hydrogen sulfide by microbial sulfate reduction (Fmsr), and hydrogen sulfide can be either buried as solid sulfide (Fis) or reoxidized (Fox). Sulfide reoxidation (biotic or abiotic) produces sulfate or an intermediate sulfur compound (elemental sulfur or sulfite) (Fox). Finally, intermediate sulfur compounds can be reoxidized back to sulfate (Fox) or further disproportionated into hydrogen sulfide (Fdh) and sulfate (Fdo) (Figure 3. 6).

The model considers the sulfur isotopes 32 S, 33 S and 34 S and calculates the sulfur isotope composition (δ 34 S, δ 33 S, ∆ 33 S) of gypsum, S° and solid sulfide formed in the marginal basin.

Model equations consist of two mass balances for each of the three boxes: one chemical mass balance for total sulfur and one isotopic mass balance for the heavy sulfur isotopes, representing 33 S or 34 S (see section 2.4 for a full model description).

Fractionation factors

Sulfur redox or phase changes are often associated with a differential partitioning of sulfur isotopes. Therefore, the expressed fractionation Ɛnet and λnet between two co-occurring sulfur Chapter 3-Cryptic sulfur cycling during the formation of giant gypsum deposits _____________________________________________________________________________________________________ 125 species of interest (typically sulfate and sulfide) results from the net effect of multiple processes involved in the cycling of these sulfur species (Figure 3. 5). To model the biogeochemical processes at play during the genesis of our samples, we assigned fixed fractionation factors established in pure culture experiments or environmental studies to each of the fractionating process considered in our model (see section 2.4.2.1), with the exception of sulfate reduction that potentially expresses a large range of isotopic fractionation [START_REF] Gomes | Sulfur isotope fractionation in modern euxinic systems: implications for paleoenvironmental reconstructions of paired sulfatesulfide isotope records[END_REF]Sim et al., 2011b;[START_REF] Wing | Intracellular metabolite levels shape sulfur isotope fractionation during microbial sulfate respiration[END_REF])(see section 2.4.2.2). In the following paragraph, we detail the approach we chose to simulate fractionation of multiple sulfur isotopes during MSR.

Laboratory experiments and studies of natural environments suggest that there is a first-order control of the specific cellular rate of MSR (csSRR) on the extent of sulfur isotopic fractionation during MSR, with fractionation approaching equilibrium values (i.e. Ɛeq =-70‰ and λeq=0.515) when csSRR reaches minimum values (Sim et al., 2011a;Sim et al., 2011b;[START_REF] Stam | 34S/32S fractionation by sulfate-reducing microbial communities in estuarine sediments[END_REF] (Figure 3. 5). The laboratory-observed csSRR at which these relationships have been established are often 1-3 orders of magnitude higher than the csSRR usually reported from natural marine environments [START_REF] Colangelo-Lillis | Diversity decoupled from sulfur isotope fractionation in a sulfatereducing microbial community[END_REF]Jørgensen, 2021 and references therein). Some authors therefore argue that csSRR in marine environments are so low that cells always express fractionation factors close to equilibrium values (ƐMSR≈-70‰ and λMSR≈0.515) (e.g. [START_REF] Masterson | Interpreting multiple sulfur isotope signals in modern anoxic sediments using a full diagenetic model (California-Mexico margin: Alfonso Basin)[END_REF]. These authors propose that when apparent fractionation factors between sulfate and sulfide (Ɛnet and λnet) deviate from equilibrium values, the deviation is mainly a consequence of sulfide pooling produced by sedimentary or early diagenetic processes (see section 3.5.2.2.1 for further explanations on this mechanism) [START_REF] Masterson | Interpreting multiple sulfur isotope signals in modern anoxic sediments using a full diagenetic model (California-Mexico margin: Alfonso Basin)[END_REF].

Chapter 3-Cryptic sulfur cycling during the formation of giant gypsum deposits _____________________________________________________________________________________________________ 126 Ɛnet>-70‰ and λnet<0.515 are often observed in systems where depositional controls are absent or negligible, implying that natural populations of microbial sulfate reducers can produce a large range of isotopic fractionations [START_REF] Colangelo-Lillis | Diversity decoupled from sulfur isotope fractionation in a sulfatereducing microbial community[END_REF][START_REF] Gilhooly | A comprehensive sulfur and oxygen isotope study of sulfur cycling in a shallow, hyper-euxinic meromictic lake[END_REF]Gomes and Hurtgen, 2015 and references therein;litterature review in Sim et al., 2011a). In addition, modeling studies also suggest a first order control of csSRR on fractionation during MSR [START_REF] Bradley | Patterns of sulfur isotope fractionation during microbial sulfate reduction[END_REF][START_REF] Wing | Intracellular metabolite levels shape sulfur isotope fractionation during microbial sulfate respiration[END_REF]. Howbeit, secondary factors independent from csSRR might also influence the extent of isotopic fractionation expressed during MSR (e.g. [START_REF] Bradley | Patterns of sulfur isotope fractionation during microbial sulfate reduction[END_REF][START_REF] Brüchert | Controls on stable sulfur isotope fractionation during bacterial sulfate reduction in Arctic sediments[END_REF][START_REF] Colangelo-Lillis | Diversity decoupled from sulfur isotope fractionation in a sulfatereducing microbial community[END_REF][START_REF] Stam | 34S/32S fractionation by sulfate-reducing microbial communities in estuarine sediments[END_REF][START_REF] Wing | Intracellular metabolite levels shape sulfur isotope fractionation during microbial sulfate respiration[END_REF]. Although it is not within the scope of this work to reconcile laboratory experiments with field observations and modelling studies, we point out that the correct estimation of csSRR in natural environments (in mol.cell -1 .day -1 ) relies on: (i) a precise estimate of the number of physiologically active sulfate reducing microbial cells, and (ii) the absence of significative reoxidative sulfur cycling. Any overestimation of active MSR cell abundance [START_REF] Colangelo-Lillis | Diversity decoupled from sulfur isotope fractionation in a sulfatereducing microbial community[END_REF] and/or underestimation of the rate of sulfide oxidation [START_REF] Jørgensen | Sulfur biogeochemical cycle of marine sediments[END_REF][START_REF] Treude | Rapid sulfur cycling in sediments from the Peruvian oxygen minimum zone featuring simultaneous sulfate reduction and sulfide oxidation[END_REF] might lead to a large underestimation of csSRR. It is possible that the current observed range of csSRR in natural environments is biased towards low rates [START_REF] Jørgensen | Sulfur biogeochemical cycle of marine sediments[END_REF], whereas csSRR is directly measured in pure culture experiments that allow the entire range of possible MSRinduced fractionations to be explored (Sim et al., 2011a;Sim et al., 2011b) (Figure 3. 5).

We consider that pure culture experiments linking fractionation factors ƐMSR and λMSR to csSRR provide a valid first order approximation of the behavior of sulfur isotope fractionation to MSR.

We used a nonlinear regression model in combination with pure culture experiment ƐMSR , λMSR and csSRR data to model the full range of MSR fractionation potentially expressed in Chapter 3-Cryptic sulfur cycling during the formation of giant gypsum deposits _____________________________________________________________________________________________________ 127 natural environments (see section 2.4.2.2 for detailed explanation). This approach does not consider the other environmental controls on MSR induced isotopic fractionation and should not be used to directly deduce a csSRR from our data but provide a simplistic way to model the ƐMSR and λMSR co-variation produced by MSR.

Results and discussion

Iron sulfides are found in all investigated gypsum and marl samples from VdG (Figure 3. 1A, supplementary section 3.8.1.1, and previously by [START_REF] Panieri | Microbial communities in Messinian evaporite deposits of the Vena del Gesso (northern Apennines, Italy)[END_REF]Sinninghe Damsté et al. (1995)). Iron sulfides form from the reaction of hydrogen sulfide produced by MSR with reactive iron, thus providing evidence for an active biogeochemical sulfur cycle during the first phase of the MSC.

Marl interbeds

In marls, iron sulfides display the smallest variability, the lowest δ 34 SIS (average -32 ±5 ‰ vs CDT (1σ)) and the highest Δ 33 SIS (average 0.12±0.04‰ (1σ)) of all analyzed samples (Figure 3. 1,Figure 3. 4,supplementary section 3.8.1.1). These isotopic values are typical of environments where the amount of available dissolved sulfate for MSR is effectively infinite (e.g. [START_REF] Gomes | Sulfur isotope fractionation in modern euxinic systems: implications for paleoenvironmental reconstructions of paired sulfatesulfide isotope records[END_REF][START_REF] Johnston | Sulphur isotopes and the search for life: strategies for identifying sulphur metabolisms in the rock record and beyond[END_REF]. We calculated net isotope fractionation factors 34 ɛnet and 33 λnet for marl samples considering a sulfate isotopic composition equal to that of the Messinian Ocean, based on the assumption that sulfide in marls was produced via the reduction of oceanic sulfate.

Marls fractionation factors fall within the MSR field (Figure 3. 5) and in fact represent the set of samples in this study that are the closest to the maximal fractionation expressed during Chapter 3-Cryptic sulfur cycling during the formation of giant gypsum deposits _____________________________________________________________________________________________________ 128 MSR, with similar values to what is observed in modern organic-rich hypoxic to euxinic basins (e.g. [START_REF] Gomes | Sulfur isotope fractionation in modern euxinic systems: implications for paleoenvironmental reconstructions of paired sulfatesulfide isotope records[END_REF][START_REF] Johnston | Sulphur isotopes and the search for life: strategies for identifying sulphur metabolisms in the rock record and beyond[END_REF]. This is consistent with previous observations suggesting that these organic rich marls formed under stratified and frequently euxinic water columns (Dela Pierre et al., 2014;Isaji et al., 2019;Lugli et al., 2010;Sinninghe Damsté et al., 1995) (Fig. 7a). (Natalicchio et al., 2022;[START_REF] Panieri | Microbial communities in Messinian evaporite deposits of the Vena del Gesso (northern Apennines, Italy)[END_REF]Schopf et al., 2012) and PLG deposits of other Mediterranean marginal basins (Dela Pierre et al., 2015;Natalicchio et al., 2022). The origin of the filaments has been disputed. They have been interpreted as cyanobacteria [START_REF] Panieri | Microbial communities in Messinian evaporite deposits of the Vena del Gesso (northern Apennines, Italy)[END_REF], even if they are generally larger (~70µm diameter) than most modern cyanobacterial filaments (Schopf et al., 2012 and references therein) ((Figure 3. 2E-I). Moreover, the fact that filaments are mostly located both in the reentrant angle and along the vertical growth bands ((Figure 3. 2D-F) suggests that they formed a benthic mat located at the gypsum-bottom water interface and that filamentous microorganisms were entrapped within gypsum in their living position, rather than after having settled from the photic zone to the basin floor (Dela Pierre et al., 2015;Schopf et al., 2012). Finally, the presence of intracellular iron sulfide globules ((Figure 3. 2G-I) -interpreted as resulting from diagenetic transformation of sulfur globules (Dela Pierre et al., 2015;Schopf et al., 2012) -is a diagnostic feature of benthic giant CSB. Such observations confirm the affiliation of the observed filament networks to chemotrophic microbial mats where both MSR Chapter 3-Cryptic sulfur cycling during the formation of giant gypsum deposits _____________________________________________________________________________________________________ 129 and sulfide oxidation were active (Dela Pierre et al., 2015;Natalicchio et al., 2022;Schopf et al., 2012).

Gypsum beds

3.5.2.2 Isotopic composition of PLG gypsum deposits 3.5.2.2.1 Influence of syn-depositionnal or early diagenetic processes on multiple sulfur isotopes The sulfur isotopic composition of the dissolved sulfate ion in equilibrium with the VdG gypsum ranged between 20.73 and 24.11 ‰ for δ 34 SSO4, and between -0.001 and 0.049 ‰ for Δ 33 SSO4 (Fig. 1A,3,4 Supplementary Table 3.2). These values fall within a narrow range that is centered around the isotopic composition of the Messinian Ocean dissolved sulfate ion (δ 34 SMESS=22.2±0.2 ‰ and Δ 33 SMESS=0.035 ±0.013 ‰) (Masterson et al., 2016) (Figure 3. 1A, Figure 3. 3 Figure 3. 4). In contrast to the almost invariant isotopic composition of the sulfate, the iron sulfides entrapped in the same gypsum samples display a large range of isotopic compositions (-36 < δ 34 SIS < +9‰ and -0.017 < Δ 33 SIS < 0.149), from high Δ 33 SIS -low δ 34 SIS, which is comparable to that of iron sulfides contained within marl deposits intercalated to the gypsum beds, to close to zero Δ 33 SIS and δ 34 SIS (Figure 3. 1, Figure 3. 4 and Supplementary section 3.8.1.1).

The theoretical maximal sulfur isotopic fractionation induced by MSR (ƐMSR≈-70‰) generate sulfide isotope values of ~-50‰ in δ 34 SIS when the sulfate used during MSR has a marine isotopic composition (~+20‰). In some cases, the isotopic composition of iron sulfide differs substantially from this maximal estimate due to sulfide pooling controlled by local diagenetic or sedimentological processes (e.g. [START_REF] Masterson | Interpreting multiple sulfur isotope signals in modern anoxic sediments using a full diagenetic model (California-Mexico margin: Alfonso Basin)[END_REF]Pasquier et al., 2021b;Pasquier et al., 2017). Sulfide pooling occurs when the rate at which sulfate is consumed (and sulfide is produced) in a given environment overcomes the rate at which the sulfate pool is replenished by diffusion from the overlying sulfate-rich water column [START_REF] Fike | Rethinking the ancient sulfur cycle[END_REF]. This leads to the progressive enrichment in 34 S of the residual sulfate and produces sulfide pools following a Rayleigh-type distillation. The observed isotopic difference (Ɛnet) between oceanic sulfate values and sedimentary sulfide is therefore smaller than the actual MSR induced fractionation (ƐMSR) because the pool of sulfate used by MSR is not oceanic sulfate Chapter 3-Cryptic sulfur cycling during the formation of giant gypsum deposits _____________________________________________________________________________________________________ 131 (and therefore replenishable) but an aliquot of it (and therefore finite). For example, sulfide pooling has been proven to control the isotopic composition of sulfides in cases where sedimentation rates are so high that porewater sulfate is disconnected from the overlying water column (Pasquier et al., 2017), or where high concentrations or reactivities of organic matter in sediments results in efficient sulfate depletion due to MSR (Pasquier et al., 2021b).

In other cases, iron sulfides enriched in 34 S are observed due to sediment mixing by oceanic currents or bioturbation bringing together pooled 34 S-enriched sulfide with newly formed 34 Sdepleted sulfide [START_REF] Fike | Rethinking the ancient sulfur cycle[END_REF]Pasquier et al., 2021a).

The petrographic evidence and isotopic signature of gypsum and co-occurring iron sulfides from the VdG deposits infers particular depositional settings that exclude the role of sulfide pooling in shaping the sulfur isotopic signatures measured in this study. We detail the reasoning in the following paragraphs.

In the δ 34 S vs Δ 33 S space, part of the gypsum iron sulfide samples fall along a curve (see the curved arrow in Figure 3. 4) consistent with the mixing of two sulfur reservoirs (Ono et al., 2006): (i) sulfide formed by MSR from an infinite, undistilled sulfate reservoir where the isotopic fractionation between sulfate and sulfide is close to equilibrium (corresponding here to the most depleted sulfide measured in this study with δ 34 S ~-40‰ and positive Δ 33 S ~+0.17‰) and (ii) what would seem to be pooled sulfide produced by the consumption of SO4 by MSR (and therefore characterized by an isotope composition that tends to approach that of initial, unreacted sulfate, as sulfate is consumed) (Ono et al., 2006;[START_REF] Zhang | Redox chemistry changes in the Panthalassic Ocean linked to the end-Permian mass extinction and delayed Early Triassic biotic recovery[END_REF].

To test the possible influence of sulfide pooling on the isotopic composition of our samples, we applied a Rayleigh type distillation to estimate the relative amount of sulfate necessary to be consumed to reproduce our data. We calculated that 55 to 95% of the sulfate reservoir should be consumed by MSR to reproduce the 34 S-enrichment in our samples in which δ 34 SIS Chapter 3-Cryptic sulfur cycling during the formation of giant gypsum deposits _____________________________________________________________________________________________________ 132 falls between -20‰ and +10‰ (capturing 50% of the iron sulfides measured) (Figure 3. 4). We tested this with fractionation factors of -70‰ (i.e., the thermodynamic equilibrium fractionation between sulfate and hydrogen sulfide) and -58‰ (i.e., the net fractionation between the lowest measured δ 34 SIS and its corresponding δ 34 SSO4).

Such significant consumption of sulfate should either: (i) have led to the undersaturation of bottom waters with respect to gypsum, thus resulting in partial dissolution of the underlying gypsum, or (ii) have produced a substantial 34 S-enrichment of the isotopic signature of the residual sulfate pool preserved in gypsum. However, detailed petrographic studies of VdG gypsum here and in both Reghizzi et al. (2018) and [START_REF] Panieri | Microbial communities in Messinian evaporite deposits of the Vena del Gesso (northern Apennines, Italy)[END_REF] indicate that gypsum laminae have grown continuously without being interrupted by dissolution episodes (Figure 3. 2) and no substantial 34 S-enrichment of the gypsum phase (Figure 3. 1,Figure 3. 3 Figure 3. 4) have been observed in our samples, consistently with Lugli et al. (2007).

Finally, a drastic drop in the sulfate concentration within CSB mats seems highly unlikely because (i) mats hosting CSB grew on gypsum lamina (Figure 3. 2) that would provide an effectively infinite source of sulfate upon gypsum undersaturation; however, no inter-laminae dissolution features are observed as stated above, and (ii) CSB can efficiently oxidize reduced sulfur to SO4 [START_REF] Brüchert | Biogeochemical and Physical Control on Shelf Anoxia and Water Column Hydrogen Sulphide in the Benguela Coastal Upwelling System off Namibia[END_REF]Callbeck et al., 2021;[START_REF] Treude | Rapid sulfur cycling in sediments from the Peruvian oxygen minimum zone featuring simultaneous sulfate reduction and sulfide oxidation[END_REF].

The combined action of biogeochemical production of SO4 by CSB, together with an infinite SO4 reservoir (represented by the gypsum substrate below the mat and the water column above the mat), exclude the possibility of a large reservoir effect due to sulfate consumption.

We therefore conclude that pooling effects did not modify the isotopic composition of our samples.

Chapter 3-Cryptic sulfur cycling during the formation of giant gypsum deposits _____________________________________________________________________________________________________ 133

Sulfide and sulfate mineral formation from the same sulfur pool

An alternative scenario is therefore needed to explain the large range of variation in δ 34 S isotopic composition of iron sulfides entrapped in gypsum (-58‰<Ɛnet<-14‰). Previous nano-SIMS studies have shown that the isotopic composition of sulfides measured within microbial mats varied with depth due to variable sulfate concentration and/or csSRR within the mat [START_REF] Fike | Micron-scale mapping of sulfur cycling across the oxycline of a cyanobacterial mat: a paired nanoSIMS and CARD-FISH approach[END_REF]. The isotopic compositions reported here, however, represent bulk measurements for each PLG gypsum cycle, and therefore are reflective of the most prominent processes governing the partitioning of sulfur species in the mat within several "generations" of CSB hosting mat (rather than local intra-mat conditions).

In the PLG deposits studied here, gypsum is a bottom-grown precipitate and is composed of an alternation of clear and turbid laminae (Figure 3. 2D-F). The mm-thick turbid laminae entrap filaments of CSB in close association with marine snow flocules rich in organic matter and iron sulfide minerals, (Figure 3. 2D-L). This suggests that sulfate reduction and sulfide reoxidation by CSB bacteria occurred in the same space, most probably in the mat adhering to the previously formed gypsum laminae (Figure 3. 2D-F, Figure 3. 8B,C). Since the surface underlying each successive layer of microbial mat was gypsum, the bottom boundary of the MSR zone was likely restricted to the base of the microbial mat, and sulfides were produced in-situ with no contribution of hydrogen sulfide diffusing from below (Figure 3. 8).

Accordingly, iron sulfides co-precipitated with gypsum have multiple sulfur isotopic values (Δ 33 SIS -δ 34 SIS) that neither reflect the overall water column conditions of the basin nor the bottom water sulfur cycling, but instead provide information on the dynamic of sulfur cycling within the mats permineralized in gypsum.

We interpret the Δ 33 SIS -δ 34 SIS signature of iron sulfides as reflecting isotope fractionation processes associated with the biogeochemical sulfur cycle during gypsum precipitation (Figure

Paired sulfate-sulfide isotopes as a fingerprint of reoxidative sulfur cycling

The combination of iron sulfide and gypsum co-precipitation and the absence of sulfide pooling influence here allows for the calculation of net fractionation factors ( 34 ɛnet -33 λnet) between sulfate from gypsum and their coeval iron sulfides in our samples.

These net sulfate-sulfide isotope fractionation factors fall between the range of isotopic fractionations observed in pure cultures of microbial sulfate reducers and those observed in pure cultures of sulfur disproportionators (Figure 3. 5). Therefore, we conclude that a process characterized by 33 λ higher than maximal 33 λMSR (>0.515) contributed, together with MSR, to define the isotopic composition of at least a part of the gypsum samples. The candidate processes include biotic H2S oxidation ( 33 λH2S-SO4ox= 0.513-0.529) and intermediate oxidation state sulfur disproportionation ( 33 λH2S-SO4disp = 0.514-0.528) (see Table 2.1). In support of this idea, the ten most 34 S-enriched samples (δ 34 SIS between -20 and +8 ‰ and Δ 33 SIS between 0.01 and 0.05‰) deviate from the mixing curve (Figure 3. 4) that assumes sulfide originating from a reservoir affected by sulfide pooling. Indeed, the contribution from a pooled sulfide reservoir is highly unlikely (see section 3.5.2.2.1), implying that this deviation is due to an additional sulfide pool. This alternative sulfide pool has a higher Δ 33 SIS than seawater (0.035‰), and was possibly produced by disproportionation, as proposed in studies where a similar deviation was observed [START_REF] Zhang | Redox chemistry changes in the Panthalassic Ocean linked to the end-Permian mass extinction and delayed Early Triassic biotic recovery[END_REF]. 34 ɛnet and 33 λnet calculated for the gypsum (massive, banded, and branching selenite) and marl samples using the isotope compositions of the sulfate ion in equilibrium with gypsum and the sulfur isotope composition of iron sulfides (dots and crosses) compared with ranges of 34 ɛ and 33 λ values obtained in pure cultures of sulfur metabolizing microorganisms (shaded areas). For marls samples, where no gypsum precipitates occur, an isotope composition of the sulfate ion equal to the Messinian ocean value (Masterson et al., 2016) was set, based on the assumption that sulfides from marls were produced through the reduction of oceanic sulfate. See figure 1.20 for references on data from abiotic sulfide oxidation experiments (dark grey area with black dotted contour), data from pure cultures with sulfate reducers (light grey area), data from pure culture experiments with sulfur disproportionators and 34 ɛnet and 33 λnet fractionation factors measured in natural environments populated with biofilms dominated by chemolithotrophic sulfide oxidizers (white squares).

This scenario is fully supported by the occurrence of two distinct sulfide pools in gypsum (Figure 3. 2Figure 3. 4): the first one, associated with marine snow floccules (Figure 3. 2J-L), was produced by MSR in an infinite sulfate reservoir; the second, associated with remains of sulfide oxidizing bacteria (Figure 3. 2G-I), results from the reoxidative part of the sulfur cycle. In support of the significant occurrence of reoxidative sulfur cycling, the sulfur and oxygen isotopic composition of the dissolved sulfate ion in equilibrium with the VdG gypsum (20.73 < δ 34 SSO4 ‰ vs CDT < 24.11; -0.001 < Δ 33 SSO4 ‰ vs CDT < 0.049‰; 8.01 < δ 18 OSO4 ‰ vs SMOW < 10.84) (Figure 3. 1Figure 3. 3Figure 3. 4 and Table 2 ) fall within a restricted range centered around the isotopic compositions of the dissolved sulfate ion of the Messinian Ocean (δ 34 SMESS=22.2±0.2 ‰ and Δ 33 SMESS=0.035 ±0.013 ‰) (Masterson et al., 2016) and δ 18 OMESS=9±2‰ (Markovic et al., 2016)), implying negligible contributions from freshwater sulfate (Figure 3. 1Figure 3. 3Figure 3. 4). Together with the very low iron sulfide concentration in gypsum (supplementary Figure 3. 9), the lack of 34 S and 18 O enrichment of sulfate characteristic of MSR [START_REF] Kah | Geochemistry of a 1.2 Ga carbonateevaporite succession, northern Baffin and Bylot Islands: implications for Mesoproterozoic marine evolution[END_REF][START_REF] Melegy | Microstructure and geochemistry studies on Messinian gypsum deposits from the Northern Coast of Egypt[END_REF] implies either that MSR activity during gypsum deposition was limited, or that the re-oxidation of sulfide into sulfate was nearly complete, i.e. the sulfur cycle was cryptic (Aller et al., 2010;Blonder et al., 2017) (Figure 3. 3).

Chapter 3-Cryptic sulfur cycling during the formation of giant gypsum deposits _____________________________________________________________________________________________________ 137 3.5.3 Modelling multiple sulfur isotopes in the Vena del Gesso basin.

Modeling strategy

To evaluate the importance of microbially-induced sulfide oxidation and to test the hypothesis of cryptic sulfur cycling, we quantified the relative contribution of biogeochemical processes in defining the multiple S isotope compositions of the VdG samples applying the numerical model described in section 3.4.3 (Figure 3. 6). In order to find a best fit to our data we tested several scenarios of sulfur redox cycling (see supplementary section 3.8.2.4), each one characterized by different relative sulfur fluxes (controlled by dimensionless fx parameters defined in section 2.4.3) and by a combination of distinct isotope fractionations (see Table 2.1) according to the following scenarios: In this best fit model run, 90% of the isotopic data in both diagrams (δ 34 SSO4 -Δ 33 SSO4, Figure 3.

7A; 34 ɛnet -33 λnet, Figure 3. 7B) plot within the area defined by 0.85 < fox < 1. The combined sulfate and sulfide isotope dataset can therefore only be reproduced by considering that: (1) at least 80% of the net incoming sulfate is reduced microbially to H2S rather than directly precipitated as gypsum (fmsr>0.8) (Fig. 8), and (2) nearly all of the MSR-produced H2S is reoxidized (0.85<fox<1) (Figure 3. 7).We suggest that the relatively high ƐMSR and fox values inferred for gypsum compared to lower values for marls indicates the occurrence of alternations between drastically different biogeochemical environments associated with periods of marl and gypsum deposition (Figure 3. 7,Figure 3. 8).

These biogeochemical changes were controlled by orbitally driven climatic oscillations triggering changes in hydrological regime. During marl deposition (i.e. precession minima), freshwater inputs were higher, leading to water column stratification, less re-oxidation of hydrogen sulfide and an efficient preservation of organic matter (Natalicchio et al., 2019;[START_REF] Sabino | The response of water column and sedimentary environments to the advent of the Messinian salinity crisis: insights from an onshore deep-water section (Govone, NW Italy)[END_REF]) (Figure 3. 1A andFigure 3. 8A). Conversely, during gypsum precipitation (precession maxima), intense reoxidation of hydrogen sulfide prevailed (Figure 3. 8B), requiring at least seasonal mixing of the water column. The large majority of sulfur (>80%)

forming gypsum was first reduced from sulfate to sulfide by MSR, possibly stored as elemental

Implications for carbon cycling and gypsum formation

To our knowledge, this is the first time that gypsum deposits are proven to be formed from sulfate sourced almost entirely from cryptic sulfur cycling. The average rate of vertical gypsum accumulation at VdG was 1.4 mm.year -1 (Reghizzi et al., 2018). Given that our model results Chapter 3-Cryptic sulfur cycling during the formation of giant gypsum deposits _____________________________________________________________________________________________________ 142 imply that more than 80% of sulfate constituting this gypsum was first reduced by MSR, we calculated that approximately 1.9 mmol.cm -2 .year -1 of sulfate was consumed by MSR. Then, by considering a rough C:S stoichiometry of 2:1 for the MSR reaction [START_REF] Jørgensen | Sulfur biogeochemical cycle of marine sediments[END_REF], we estimated a gross benthic organic carbon flux of 3.8 mmol.cm -2 .year -1 remineralized by MSR during PLG gypsum formation. This is consistent with the large amounts of organic matter contained in the VdG interbeds. Marl deposits have a total organic carbon (TOC) content comparable to that of euxinic environments [START_REF] Jovovic | Quantitative and specific recovery of natural organic and mineral sulfur for (multi-)isotope analysis[END_REF]Lugli et al., 2007;Sinninghe Damsté et al., 1995), whilst TOC concentrations in gypsum range between 0.1 to 0.96% (Lugli et al., 2007), which is substantial considering the enormous dilution effect arising from gypsum depositional rates that are 25 to 250 times greater than marls (Reghizzi et al., 2018)).

The benthic carbon flux involved in MSR we estimate for the VdG basin is comparable to that occurring in the world's most productive ocean regions, such as the Benguela and Peru-Chile margins, where the organic carbon remineralized through MSR ranges between 0.2 and 20mmol.cm -2 .yrs -1 [START_REF] Bowles | Global rates of marine sulfate reduction and implications for sub-sea-floor metabolic activities[END_REF][START_REF] Treude | Rapid sulfur cycling in sediments from the Peruvian oxygen minimum zone featuring simultaneous sulfate reduction and sulfide oxidation[END_REF]. In these areas, dense mats populated by microbial sulfate reducers and CSB, similar to the fossil microbial assemblages found in the VdG gypsum, fuel a cryptic sulfur cycle [START_REF] Brüchert | Biogeochemical and Physical Control on Shelf Anoxia and Water Column Hydrogen Sulphide in the Benguela Coastal Upwelling System off Namibia[END_REF]Callbeck et al., 2021;[START_REF] Treude | Rapid sulfur cycling in sediments from the Peruvian oxygen minimum zone featuring simultaneous sulfate reduction and sulfide oxidation[END_REF].

The above considerations do not necessarily imply that the VdG basin was characterized by similar primary productivity to modern oxygen minimum zones but, rather, by comparable benthic carbon fluxes and carbon and sulfur cycling rates. Indeed, aerobic consumption of organic carbon in the VdG basin may have been limited to the upper layer of the water column due to density stratification (Natalicchio et al., 2022), leading to high organic carbon export fluxes (Pellegrino et al., 2021) and large quantities of benthic organic carbon available for MSR Chapter 3-Cryptic sulfur cycling during the formation of giant gypsum deposits _____________________________________________________________________________________________________ 143 (Figure 3. 8). Conversely, very high primary productivity controls the export flux of organic carbon in modern productive margins.

Our findings support a recent model of PLG deposition [START_REF] Aloisi | The geochemical riddle of "low-salinity gypsum" deposits[END_REF] which suggests that the activity of benthic CSB mats drove increases in dissolved sulfate concentration and triggering gypsum precipitation. The inhibition of gypsum precipitation in deeper, more distal lateral equivalents of PLG gypsum (Natalicchio et al., 2019;[START_REF] Sabino | The response of water column and sedimentary environments to the advent of the Messinian salinity crisis: insights from an onshore deep-water section (Govone, NW Italy)[END_REF], together with our results showing that more than 80% of the sulfate constituting the more marginal VdG gypsum has been cycled through a dynamic sulfur cycle, suggest that gypsum precipitation could have been controlled by the efficiency of the re-oxidative portion of the sulfur cycle, i.e. the cryptic cycling of sulfur.

Concluding remarks

Our work suggests a tight coupling of microbially driven sulfate reduction and sulfide oxidation in the Vena del Gesso basin.

Primary Lower Gypsum deposits accumulated during the first phase of the Messinian Salinity Crisis and characterized by identical lithofacies, stacking patterns, petrographic features and microbial fossils are widespread across the Mediterranean (see Figure 3. 1C) (Lugli et al., 2010;Natalicchio et al., 2022;Pellegrino et al., 2021). Whereas these deposits were previously thought to be restricted to semi-enclosed basins with a water column depth of no more than 200m (Lugli et al., 2010), recent studies have demonstrated that some were deposited at depths of up to 800m and in basins not necessarily restricted by a sill (Ochoa et al., 2015;Raad et al., 2021). If the cryptic sulfur cycle we describe was also active during the deposition of gypsum elsewhere in the Mediterranean, this would imply that the restriction of the Atlantic connection at the end of the Miocene led to not only the widespread deposition of evaporites, Chapter 3-Cryptic sulfur cycling during the formation of giant gypsum deposits _____________________________________________________________________________________________________ 144 but also the significantly altered the short-term carbon and sulfur cycles of the Mediterranean [START_REF] Waldeck | Calibrating the triple oxygen isotope composition of evaporite minerals as a proxy for marine sulfate[END_REF]. Further, the extensive occurrence of microbial mats populated by CSB on modern high-productivity continental margins or restricted basins, together with the results of the present study, suggest that cryptic sulfur cycling might have been prevalent in analogue evaporitic environments of the geological past. Iron sulfide concentrations calculated from our chromium reducible compound chemical extractions range from 15 to 502 ppm in selenite samples and from 1155 to 13117 ppm in marls (Table 3. 1). The higher concentration of iron sulfide in marly interbeds are comparable to sulfide concentration found in typical ancient and modern euxinic environments [START_REF] Jovovic | Quantitative and specific recovery of natural organic and mineral sulfur for (multi-)isotope analysis[END_REF]. Instead, sulfide concentrations in gypsum samples is much lower and variable (Figure 3. 9).

Chapter 3-Cryptic sulfur cycling during the formation of giant gypsum deposits _____________________________________________________________________________________________________ 148 The isotopic composition of riverine input was set equal to the average values of modern rivers (δ 34 SSO4=4.8‰ (Burke et al., 2018) and Δ 33 SSO4=-0.0187‰ (Tostevin et al., 2014)). Based on the observation by Topper et al. (2011) that riverine inputs contributed for 10 to 80% of the total water inputs during PLG formation and on the fact that freshwater inputs are approximately 100 times less concentrated than oceanic waters in sulfate ion (Burke et al., 2018), we calculated that the fraction of incoming sulfate originating from riverine inputs on the total sulfate input, friv, ranged between 0 and 0.075, we fixed it here at 0.035. This is consistent with the blue mixing curve displayed on Figure 3. 4 that describe the mixing between oceanic sulfate and riverine inputs and from which can be inferred that the riverine inputs had a negligible influence on the sulfate isotopic composition of gypsum.

We then fixed the external basin sulfate isotopic composition equal to that of the sulfate from the first gypsum bed of the Vena del Gesso basin (δ 34 SSO4=21.78‰ and Δ 33 SSO4=0.033‰). This isotope composition is very close to that of sulfate dissolved in the global Messinian ocean (δ 34 SMESS=22.16±0.3 ‰ and Δ 33 SMESS= = 0.035 ±0.013 ‰ (Masterson et al., 2016)). Hydrochemical models of the MSC suggest that during deposition of PLG a two-way (inflow-outflow) water exchange always existed between the Mediterranean Sea and the Atlantic Ocean, and between the marginal basins and the main Mediterranean basin (Topper et al., 2014). At the acme of gypsum precipitation in a marginal basin connected to the Mediterranean sea, [START_REF] Topper | The precessional phase lag of Messinian gypsum deposition in Mediterranean marginal basins[END_REF] calculated that calcium sulfate concentration in the marginal basin is approximately 1.2 times greater than in the main Mediterranean basin, whereas the water outflow represents 69 to 77% of the water inflow from the main Mediterranean basin. This allowed us to estimate that the fraction of sulfate output from the Vena del Gesso basin to Chapter 3-Cryptic sulfur cycling during the formation of giant gypsum deposits _____________________________________________________________________________________________________ 152 the Mediterranean represented between 87 and 96% of the incoming sulfate fluxes; thus, we set fex equal to 0.9 to represent sulfate exchanges with the external sulfate source during gypsum precipitation. The amount of intermediate sulfur compound being disproportionated into sulfide fdh is the complement of fdo, the amount of intermediate sulfur compound being disproportionated into sulfate. fdh was chosen depending on the scenario tested, with fdh=0 when the model was tested for the absence of reoxidation or disproportionation, fdh=0.25 when we tested the model for disproportionation of sulfite and fdh=0.75 when the elemental sulfur disproportionation scenario was choosen (see Pellerin et al., 2015 and references therein for stoichiometry of disproportionation reactions). Finally, the entire range of values for the fraction of sulfur reduced through MSR (0 < fmsr < 1) was tested in order to find the fmsr that would best fit our data (see supplementary information). 0.85 (brown area). The main differences from our best fit results (where fex =0.9) concern the sulfate isotopic values, with the area covered being smaller when the choosen fex is larger, possibly requiring a larger fmsr to explain the whole range of values covered by our sulfate data (see supplementary section 6 for further explanation on this point). Thus, with fox and ƐMSR constrained as explained above, fmsr can be constrained considering the δ 34 S-Δ 33 S dataset (Figure 3. 11). The sulfate isotope composition of VdG samples spans a small range from 20.7 to 24.1 ‰ in δ 34 S and from -0.001 to 0.0049 in Δ 33 S. In the δ 34 S-Δ 33 S space, the size of the area corresponding to the full considered range of fox and ƐMSR depends on fmsr: small values of fmsr produce a limited isotope reservoir effect, and thus a small deviation of model results from the initial basin sulfate composition, while elevated fmsr values result in a strong isotope reservoir effect and a large area defined by the model results (Figure 3. 11).

With the constraints on fox and ƐMSR introduced above (fox > 0.85, ƐMSR >-50‰), only values of fmsr > 0.8 result in a good fit of the model to the sulfate data in the δ 34 SSO4-Δ 33 SSO4 space (Figure 3. 11).

Supplementary discussion on the reoxidative portion of the sulfur cycle

In the best fit model run ( Figure 3. 7), the use of fractionation factors and sulfur fluxes typical of elemental sulfur redox processing is in good agreement with the occurrence of iron sulfides associated with the Chapter 3-Cryptic sulfur cycling during the formation of giant gypsum deposits _____________________________________________________________________________________________________ 156 remains of CSB in our samples, interpreted to result from early diagenesis of elemental sulfur.

However, the nature of the microbial process driving the re-oxidative portion of this cryptic sulfur cycle is still uncertain. The best fit model run is based on the combination of sulfide oxidation to S° by chemolitotrophs biofilms and subsequent microbial S° disproportionation ( Figure 3. 7), because multiple sulfur isotope fractionation for these processes have been determined (section 2.4.2.1). However, the available isotope data for pure culture of sulfideoxidizer is scant and we cannot exclude that other re-oxidative processes involving the S° intermediate produce a comparable isotope fractionation. In particular, the abundance of microbial fossils indicating the presence of CSB mats in Vena del Gesso gypsum makes these microorganisms the best possible alternative. To our knowledge, there are no experimental studies of sulfur fractionation by pure cultures of CSB. However, CSB being chemolithotrophs, and the biofilms from which comes the fractionation factors we used for our best fit for the oxidation of H2S to S 0 being dominated by chemolithotrophs organisms, it is likely that CSB produce a similar isotopic fractionation than the one we used while oxidizing H2S. Note that the occurrence of a reoxidative sulfur cycling involving the fractionation factors for elemental sulfur disproportionation has been previously suggested in porewaters displaying similar higher 33 λnet than maximal λMSR and in the presence of Beggiatoa CSB (Pellerin et al., 2015). Thus, our data and modelling imply that the cryptic sulfur cycle was driven by MSR and a reoxidative process isotopically consistent with H2S oxidation to S 0 and subsequent elemental S° disproportionation, possibly involving the metabolism of CSB.

Other biogeochemical scenarios

In order to find a best fit to our data, we tested several scenarios involving different pathways for sulfur redox cycling. In all the tested scenarios the full range of fox (the relative amount of The occurrence of MSR without subsequent reoxidation was tested with attributing a fox=0, fdh=0 and varying fmsr and ƐMSR (Figure 3. 12). In Δ 33 S-δ 34 S diagram whereas our iron sulfide data does not fit the model, SO4 data could be explained by low values of ƐMSR and fmsr. In the 34 ɛnet and 33 λnet diagram, the model produces a curve that does not fit our data. Note that in the absence of reoxidation the model in the 34 ɛnet and 33 λnet is not an area but a curve following the different values of csSRR. This shows that the value of fmsr used in the 34 ɛ-33 λ space has no influence on the isotopic fractionation expressed. For each of the subsequent scenarios different values of fmsr (from 0 to 1) were tested in order to find a best fit to our data. This allowed us to understand what fraction of sulfate from the basin has been reduced through microbial sulfate reduction activity during PLG formation.

The direct reoxidation of hydrogen sulfide to sulfate was tested by combining the fractionation factors αox for the reoxidative flux fox attributed to equations 3, 4 or 5 (inorganic oxidation) or 6 (phototrophic oxidation) from section 2.4.1 to fdh=0 (implying fdo=1) and no fractionation for the disproportionation flux αdh (Figure 3. 11). In both Δ 33 S-δ 34 S and 34 ɛ and 33 λ the model does not cover the full range of values expressed in our samples, supporting our conclusions that the filamentous fossils found in gypsum does not corresponds to cyanobacteria remains [START_REF] Panieri | Microbial communities in Messinian evaporite deposits of the Vena del Gesso (northern Apennines, Italy)[END_REF] but to CSB remains (Dela Pierre et al., 2015;Natalicchio et al., 2022;Schopf et al., 2012). Note that in the model that consider phototrophic oxidation of hydrogen sulfide, low ƐMSR and fox values in the model approach the values obtained for iron sulfide in marls, which is in good agreement with periodic euxinia during marl deposition (Figure 3. 13). This support the observation by Sinninghe Damsté et al.

(1995) that have found abundant lipid biomarkers indicative of phototrophic green sulfur bacteria within the marls. We weren't able to test the hypothesis of a complete oxidation of sulfide into sulfate by chemolithotrophic bacteria because no fractionation factors are so far available in the literature.

The disproportionation of elemental sulfur into sulfide was tested by combining the fractionation factors αox for the reoxidative flux attributed to equations 6.1 (phototrophic oxidation) or 7.1 (chemolithotrophic oxidation) of section 2.4.1 to fdh=0.75 (implying fdo=0.25) (see the best fit Figure 3. 7) in agreement with the stoichiometry of the reaction (Thamdrup et al., 1993) whereas the disproportionation of sulfite was tested by using fdh=0.25 (fdo=0.75) (Bak and Cypionka, 1987) (Figure 3. 14). As the fractionation factors of paired 33 S and 34 S isotopes for the oxidation of aqueous sulfide into sulfite remains unknown, when we tested our model for Chapter 3-Cryptic sulfur cycling during the formation of giant gypsum deposits _____________________________________________________________________________________________________ 160 disproportionation of sulfite, the production of sulfite was done by testing all the fractionation factors available for aqueous sulfide oxidation (equations 3,4,5,6.1 and 7.1 in section 2.4.1).

Whereas the areas covered by this model are larger than for previous scenarios (Figure 3. 14), allowing to fit a larger set of data point with the model, this scenario still doesn't reproduce the whole range of isotopic data we measured for the Vena del Gesso section. Moreover, the occurrence of remains of sulfur globules in our samples point its greater implication compared to sulfite as an intermediate sulfur compound in the genesis of the observed isotopic signatures of our samples. At the end of the Miocene, the restriction of the Atlantic-Mediterranean seaway led to the deposition of a basin-wide salt giant (of up to 2.5km in thickness) in the Mediterranean (Haq et al., 2020). This event, the so-called Messinian salinity crisis (MSC), lasted less than 650 kyrs.

Introduction

Importantly, drawdown(s) of the water level of the Mediterranean have been proposed to have occurred during this event. However, their number, timing and amplitude are still largely debated.

The onset of the MSC is recorded in shallow (<800m), marginal, Mediterranean basins in the form of primary lower gypsum (PLG) consisting in cyclic alternation of bottom grown gypsum and marls (Krijgsman et al., 1999;Lugli et al., 2010;Ochoa et al., 2015;Raad et al., 2021). At Following salt deposition, the third and final stage of the MSC (stage 3) is characterized by the deposition of gypsum layers alternating with marls, occasionally hosting brackish shallow water fauna mostly endemic of the Eastern Paratethys region (Andreetto et al., 2021a and references therein). These stage 3 evaporitic deposits are nowadays either located onland, on top of clastic deposits or erosional surfaces (Andreetto et al., 2021a;Karakitsios et al., 2017;[START_REF] Manzi | A new facies model for the Upper Gypsum of Sicily (Italy): chronological and paleoenvironmental constraits for the Messinian salinity crisis in the Mediterranean[END_REF], or offshore, above the halite, at depths of up to 4000m (Hsü et al., 1978a;[START_REF] Hsü | Initial reports of the deep sea drilling project: DSDP volume XLII Part[END_REF][START_REF] Kastens | Proc. ODP, Init. Repts[END_REF]Ryan et al., 1973). Whereas PLG deposits are considered to have formed mainly in peripheral basins that were relatively shallow, and filled by continuous marine influxes (Lugli et al., 2010;Raad et al., 2021), the water level of Mediterranean basins during the accumulation of halite and stage 3 gypsum deposits is debated. Specifically, the occurrence of numerous evidences for erosion in both eastern [START_REF] Camerlenghi | Uncovering the Mediterranean Salt Giant (MEDSALT)-Scientific Networking as Incubator of Cross-disciplinary Research in Earth Sciences[END_REF][START_REF] Madof | Discovery of vast fluvial deposits provides evidence for drawdown during the late Miocene Messinian salinity crisis[END_REF]Micallef et al., 2018a;Micallef et al., 2018b) and western basins [START_REF] Andreetto | Freshening of the Mediterranean Salt Giant: controversies and certainties around the terminal (Upper Gypsum and Lago-Mare) phases of the Messinian salinity crisis[END_REF][START_REF] Lofi | Erosional processes and paleo-environmental changes in the Western Gulf of Lions (SW France) during the Messinian Salinity Crisis[END_REF]Lofi et al., 2011b) (Figure 4. 1) of the Mediterranean strongly suggest that a significant sea level fall accompanied the formation of stage 2 halite and/or stage 3 gypsum deposits. In addition, the facies and sedimentary structures of the uppermost part of stage 3 evaporites recovered during DSDP drilling expeditions have been first interpreted to represent deposition under shallow water (above wave base) [START_REF] Hsü | History of the Mediterranean salinity crisis[END_REF]Hsu et al., 1978;Hsü et al., 1973). However, some authors claim that the above-mentioned sedimentary and mineralogical evidence is equivocal. They claim that the laminated and clastic textures of stage 3 evaporites indicate their formation under fully subaqueous conditions (below sea-wave), implying a relatively high water level during MSC stage 3 (Hardie and Lowenstein, 2004;[START_REF] Lugli | The deep record of the Messinian salinity crisis: evidence of a non-desiccated Mediterranean Sea[END_REF]. The number, timing and amplitude of sea level drawdown events during the MSC are still strongly debated, with estimates ranging from 200m to 2km (Aloisi et al., (in prep); Andreetto et al., 2021a;[START_REF] Gvirtzman | Intra-Messinian truncation surface in the Levant Basin explained by subaqueous dissolution[END_REF][START_REF] Andreetto | Freshening of the Mediterranean Salt Giant: controversies and certainties around the terminal (Upper Gypsum and Lago-Mare) phases of the Messinian salinity crisis[END_REF]Lofi et al., 2011b;[START_REF] Madof | Discovery of vast fluvial deposits provides evidence for drawdown during the late Miocene Messinian salinity crisis[END_REF]Roveri et al., 2014c).

Further, the amplitude of water level drawdown might have been different between eastern and western Mediterranean basins [START_REF] Andreetto | Freshening of the Mediterranean Salt Giant: controversies and certainties around the terminal (Upper Gypsum and Lago-Mare) phases of the Messinian salinity crisis[END_REF]Micallef et al., 2018a).

In marine environments, water depth is a principal driver of sedimentary and biogeochemical processes, greatly influencing the type of sediments deposited (Aller, 2014;Middelburg et al., 1997).

In our specific MSC case, the magnitude of the water level drawdown (XX' in Figure 4. 2) has large implications for the reconstruction of Mediterranean paleoenvironmental settings. receiving temporally and spatially variable amounts of freshwater (e.g. [START_REF] Camerlenghi | Uncovering the Mediterranean Salt Giant (MEDSALT)-Scientific Networking as Incubator of Cross-disciplinary Research in Earth Sciences[END_REF][START_REF] Caruso | The late Messinian "Lago-Mare" event and the Zanclean Reflooding in the Mediterranean Sea: new insights from the Cuevas del Almanzora section (Vera Basin, South-Eastern Spain)[END_REF]Cita and Ryan, 1978) or (ii) formed when the Mediterranean basin was at the level of the global ocean, and fed by Atlantic sea water or by the Parathethys (or both) with the various sub-basins of the Mediterranean fully [START_REF] Carnevale | Fossil marine fishes and the "Lago Mare" event: Has the Mediterranean ever transformed into a brackish lake?[END_REF]Carnevale and Schwarzhans, 2022;García-Veigas et al., 2018;Roveri et al., 2014b;Roveri et al., 2014c;Vasiliev et al., 2017), or at least intermittently, connected to each other (Andreetto et al., 2022a;[START_REF] Rouchy | Paleoenvironmental changes at the Messinian-Pliocene boundary in the eastern Mediterranean (southern Cyprus basins): significance of the Messinian Lago-Mare[END_REF]. Whereas the latter scenario suggests drawdown(s) (XX' in The relation between thickness of the water column, crustal vertical motions and water level drawdown can be summarized as follows (Figure 4. 2):

X'Y' = XY + YY' -XX' et al., 1996;Hsü et al., 1973;[START_REF] Kastens | The geological evolution of the Tyrrhenian Sea: An introduction to the scientific results of ODP Leg 107[END_REF]. These samples consist of clastic or cumulated calcium sulfate minerals [START_REF] Lugli | The deep record of the Messinian salinity crisis: evidence of a non-desiccated Mediterranean Sea[END_REF], meaning that part of the calcium sulfate minerals sampled in deep basins might have an allochthonous origin.

Therefore, the petrographic features of deep ODP/DSDP evaporites have to be studied carefully before extracting hydrological or paleoenvironmental information from their strontium or sulfate isotopic signatures (Andreetto et al., 2021a;Roveri et al., 2014b).

Chapter 4-Late Messinian Mediterranean base level fall ___________________________________________________________________________________________ 167

A potential tool to provide further information on their initial depositional settings are iron sulfides, that represent minor components of the layers preserved in IODP/DSDP cores.

As opposed to gypsum, fine grained micrometric to nanometric iron sulfides are ubiquitous in marine environments and do not only form in evaporitic basins. Further, they rapidly oxidize when eroded or transported (Choppala et al., 2017;Luther et al., 1982;Morse, 1991). Because they are authigenic, their isotopic composition reflects the processes that were ongoing in the environment where they were formed and preserved (Marin-Carbonne et al., 2022;[START_REF] Riedinger | Sulfur Cycling in an Iron Oxide-Dominated, Dynamic Marine Depositional System: The Argentine Continental Margin[END_REF]. Hydrogen sulfide is produced by microbial sulfate reduction (MSR) and is highly reactive in aqueous environments [START_REF] Rickard | Sulfidic sediments and sedimentary rocks[END_REF]. It is then scavenged by reactive iron to form iron sulfide with negligible sulfur isotopic fractionation (Eldridge et al., 2021). Therefore, the iron sulfide isotopic signature reflects the isotope fractionation associated to the dominant biogeochemical processes responsible for sulfide formation/preservation within the zone of the sediment where both reactive iron and hydrogen sulfide are available. As we explain below, the thickness of the overlying water column exerts a strong control on the δ 34 SFeS isotope composition of sedimentary sulfides, providing a means to estimate paleo-thicknesses of the water column during their formation.

The S-isotope composition of iron sulfide is less homogeneous in time and space than that of sulfate bearing minerals. This is because δ 34 SSO4 reflects the large, well-mixed oceanic reservoir of sulfate while δ 34 SFeS reflects the highly dynamic and variable nature of the early diagenetic biogeochemical sulfur cycle within marine sediments, that involves a smaller reservoir of sulfur [START_REF] Fike | Rethinking the ancient sulfur cycle[END_REF]Houghton et al., 2022;Pasquier et al., 2021a;Pasquier et al., 2021b;Pasquier et al., 2017).

Several factors have been demonstrated to control the isotopic composition of preserved iron sulfides. These include : sedimentation rates and grain size, that control the porosity of the (Pasquier et al., 2017); organic carbon fluxes and type, that control rates of microbial sulfate consumption (Houghton et al., 2022;Mengdi et al., 2022;Pasquier et al., 2021b); and the occurrence of reoxidation either (i) abiotic due to episodic reworking of sediments (Aller et al., 2010;Pasquier et al., 2021a) or large metal oxides availability [START_REF] Antler | The Sedimentary Carbon-Sulfur-Iron Interplay -A Lesson From East Anglian Salt Marsh Sediments[END_REF]Blonder et al., 2017;[START_REF] Findlay | Sulfide oxidation affects the preservation of sulfur isotope signals[END_REF]Liu et al., 2020b;[START_REF] Mills | Geochemical evidence for cryptic sulfur cycling in salt marsh sediments[END_REF], or (ii) mediated by the presence of specific microbial communities (Magnuson et al., 2020;Pellerin et al., 2015;Zerkle et al., 2016).

These controls are strong and variable and can lead to high amplitude isotopic variations in shallow water marine environments (<1000m), whereas they generally have a lesser and constant influence in deep basins (>1000m) (Aller, 2014;Dunne et al., 2007;[START_REF] Jørgensen | Sulfur biogeochemical cycle of marine sediments[END_REF]Middelburg et al., 1997). Water column thickness (X-Y), through the variety of environmental and diagenetic controls listed above, has thus been suggested to exert a first order control on sedimentary sulfide isotopic composition [START_REF] Fike | Rethinking the ancient sulfur cycle[END_REF]Houghton et al., 2022;Pasquier et al., 2021a).

Following this idea, we performed a detailed SEM/EDX analyses of representative samples of each lithofacies in DSDP and IODP cores 134, 374, 376 and 654 and measured the isotopic composition of iron sulfide extracted from 45 samples from these same cores. We propose in this study that δ 34 S of early diagenetic iron sulfide can be used to approximately set a maximal water column thickness (X'Y') under which stage 3 deposits formed. We then used these estimates as an equivalent of the deepest bathymetry possible for the deposition of stage 3 sediments (Y'). This was used as an input in isostasy-based reconstructions of vertical crustal movements (YY'), to constrain the amplitude of sea level fall XX' necessary to fit our water thicknesses estimates (X'Y') during the accumulation of stage 3 deposits.

Messinian evaporites from DSDP/ODP drill cores

The accumulation of Stage 3 deposits lasted 200kyrs and follows the deposition of a km-thick halite sequence [START_REF] Manzi | A new facies model for the Upper Gypsum of Sicily (Italy): chronological and paleoenvironmental constraits for the Messinian salinity crisis in the Mediterranean[END_REF]van der Laan et al., 2006). Regardless of the water level in Mediterranean basin during stage 3, as the evaporites formed at that time consist of calcium sulfate minerals that precipitate at salinities lower than halite, they represent a lowering of the overall Mediterranean salinity compared to the preceding stage of the crisis when halite formed (Andreetto et al., 2021a;Roveri et al., 2014a Paratethys influxes during mudstone deposition [START_REF] Popescu | Lago Mare episodes around the Messinian-Zanclean boundary in the deep southwestern Mediterranean[END_REF], while at site 654 in the Tyrrhenian basin where stage 3 deposits lie at 2451mbsl [START_REF] Cita | Messinian Paleoenvironments: New Evidence from the Tyrrhenian Sea (ODP LEG 107)[END_REF] and site 376 at the Florence Rise where stage 3 is at 1965 mbsl [START_REF] Benson | The paleoecology of the ostracods of DSDP Leg 42A[END_REF]De Deckker et al., 1988) the occurrence of abundant cyprideis ostracods within the mudstones suggest their deposition under shallow water environments.

In the Ionian basin (site 374), stage 3 deposits have been drilled at 4460mbsl and diatom assemblages representative of a lacustrine to brackish water body support the hypothesis of the disconnection of the Ionian basin from the Mediterranean during the deposition of marl interbeds (Schrader and Gersonde, 1978).

Finally, the sulfate ions constituting gypsum layers generally have a marine signature (see isotopic compilation in Andreetto et al., 2021a;Andreetto et al., 2022a) (and Chapter 5 of this manuscript), whereas their 87 Sr/ 86 Sr signatures depart significantly from the coeval Messinian global ocean strontium isotope curve. This suggests their formation from a mixture of waters bearing oceanic and freshwaters isotopic signatures (Andreetto et al., 2021a;Andreetto et al., 2022a;Andreetto et al., 2022b;Andreetto et al., 2021b). The discrepancy between strontium and sulfate isotopic signature can be explained by the fact that riverine waters are generally at least 100 times less concentrated than seawater in sulfate (Burke et al., 2018). The δ 34 S isotopic composition of sulfate in a marine basin is thus less sensitive to freshwater influxes compared to strontium.

Whether or not the halite and low salinity fossils assemblages described above represent in place features, the rapid facies alternations between calcium sulfate salts and non-evaporitic interbeds and their isotopic signatures suggest the occurrence of water level and salinity variations in deep basins of the Mediterranean during MSC stage 3. The cyclical alternation of facies during stage 3 is believed to be controlled by precession driven climate oscillations [START_REF] Manzi | A new facies model for the Upper Gypsum of Sicily (Italy): chronological and paleoenvironmental constraits for the Messinian salinity crisis in the Mediterranean[END_REF]. Unfortunately, an age correlation between the samples included in this study is difficult to establish (Andreetto et al., 2022a). Further, IODP and DSDP cores are characterized by: (i) poor recovery of Messinian intervals, (ii) the absence of recovery of pre-stage 3 deposits in most of these cores and (iii) the occurrence of intra erosion/dissolution in some or all of the investigated sites (Figure 4. 3).

These features complicate the tuning of evaporite and marl alternations from the offshore record to astronomical reference curves. While the Stage 3 origin of the samples investigated in this study is clear because they overlie halite deposits and precede Pliocene sediments (Roveri et al., 2014b), it is impossible to know for each individual site at which point during the MSC phase 3 they were deposited.

Non-evaporitic interbeds of site 374 are charecterised by high Total Organic Carbon (TOC of 1.9 to 4.7%) (Deroo, 1978) with C/N>20 (Sigl et al.) suggesting a terrestrial origin of the organic matter and supporting the prevalence of freshwater inputs in the basin. No information on the origin of the organic matter in other intervals used in this study are available in the literature; however, the TOC content of other non-evaporitic intervals is generally relatively low for sites 376 (Sigl et al.) and 134 (McIver and RD, 1973) with values from 0.1 to 0.8%. and more variable in marl interbeds at site 654 with values from 0.03 to 3.7% (Pierre and Rouchy, 1990). No TOC analyses have been performed in evaporitic interbeds; however, the general whitish aspect of these layers suggest that low amount of organic matter was preserved within these layers. the interpretation of the paleoenvironmental settings ruling stage 3 deposits formation highly debated (Andreetto et al., 2021a). One of the only certainties regarding these deposits is that the calcium sulfate samples from DSDP/IODP cores were necessarily deposited in an environment at least saturated with respect to gypsum to allow their preservation, indicating hypersalinity of bottom waters of the basins during their accumulation in the offshore Mediterranean.

Material and Methods

Petrographic analyses

A total of 16 representative samples of laminated, clastic and porphyroblastic gypsum together with anhydrite and marls were studied with a scanning electron microscope (SEM) coupled with an energy dispersive X-ray spectrometer (EDS) (Table 4. 1 Table 4. 2). Analyses were performed at the electron microscopy platform at the Institut de Physique du Globe de Paris (IPGP) using a Carl Zeiss EVO MA10 SEM in back scattered electron and secondary electron imaging mode to characterize the reduced sulfur species of our samples. SEM imaging and EDS analyses were performed using an accelerating voltage of 15kV at a working distance of 12mm at an electron beam current of 2-3nA. Samples were coated with a few nanometers of Au prior to analyses.

Sulfur isotope analyses

Iron sulfides from gypsum and marls were reduced into H2S using a CrCl2 solution prepared daily. Solid CrCl3 was stirred for 15min with HCl and zinc granules under a N2 atmosphere in order to reduce Cr(III) into Cr(II) and produce CrCl2 solution (Fossing and Jørgensen, 1989).

Chapter 4-Late Messinian Mediterranean base level fall ___________________________________________________________________________________________ 174 Samples were reacted for approximately 2hours under N2 atmosphere to produce H2S that was transformed into Ag2S according to [START_REF] Geng | A simple and reliable method reducing sulfate to sulfide for multiple sulfur isotope analysis[END_REF]. Ag2S powder was then reacted with F2 to produce gaseous SF6, purified (cryogenically and chromatographically) and analyzed by a dual inlet mass spectrometer MAT-253 (Ono et al., 2006b;[START_REF] Labidi | Determination of multiple sulfur isotopes in glasses: A reappraisal of the MORB delta S-34[END_REF]Johnston et al., 2007). The δ 34 S values were calibrated with an

Chapter 4-Late Messinian Mediterranean base level fall ___________________________________________________________________________________________ 175 internal SF6 standard using the IAEAS-1 international standard that has a value of -0.3‰ VS-CDT [START_REF] Brand | Assessment of international reference materials for isotope-ratio analysis (IUPAC Technical Report)[END_REF] (see Table 4. 1Table 4. 2 for an inventory of the samples analysed).

The full method with references can be found in Chapter 2, section 2.3, tables with dataset are in Appendix A. In the next section we provide details on the strategy, parameters and data used in both scenario for (i) a full basin (ii) a significant drawdown of the water level. Mediterranean, Florence Rise) such a detailed reconstruction is not available, but a first order knowledge of the depositional environments before and after gypsum deposition can be assessed with the well data from the initial drilling reports (Bryant et al., 2020;Hsü et al., 1978a;[START_REF] Hsü | Initial reports of the deep sea drilling project: DSDP volume XLII Part[END_REF][START_REF] Kastens | Proc. ODP, Init. Repts[END_REF][START_REF] Nesteroff | Un modéle pour les évaporites messinienne en Méditerranée, bassins peu profonds avec dépot d'évaporites lagunaires[END_REF]Ryan et al., 1973). Where relevant, 0-D isostatic compensation was applied for the Pliocene-Quaternary sediment load to estimate the post-MSC subsidence at the site, and other plausible sources of vertical motions (compaction, thermal relaxation) were considered. For this scenario, the paleo-water depth X'Y'full was calculated using the following equation (Figure 4.4,panel B):

𝑷𝒂𝒏𝒆𝒍 𝑩: 𝑋′𝑌′ ,-.. = 𝑃𝑄 2 + 𝑃𝑄 45 * 71 -: 𝜌 < -𝜌 = 𝜌 > -𝜌 = ?@
Where PQD is the modern water depth of the top of Plio-Quaternary sediments, X'Y'full is stage 3 water depth for a full basin without drawdown, PQTH is Plio-Quaternary sediment thickness, ρs is sediment density, ρw is water density, and ρa is asthenosphere density.

All thicknesses/depths are in units of meters, and densities in kg/m 3 . For simplicity we apply the same density for sediment, water and the asthenosphere as reported in Heida et al.

(2021). The uncertainty introduced by variable density is much smaller than that introduced Where X'Y'max is the reference maximum water depth deduced from our isotopic investigations. And X' is the water level obtained after the drawdown and the subsequent isostatic rebound. dolomite, quartz or clay minerals that represent the major constituents of our samples [START_REF] Borsetti | Messinian and pre-Messinian sediments from ODP Leg 107 Sites 652 and 654 in the Tyrrhenian Sea: sedimentologic and petrographic study and possible comparisons with Italian sequences[END_REF][START_REF] Hsü | History of the Mediterranean salinity crisis[END_REF][START_REF] Nesteroff | Un modéle pour les évaporites messinienne en Méditerranée, bassins peu profonds avec dépot d'évaporites lagunaires[END_REF]

Results

Petrography

Sulfur isotope composition

The isotope composition of iron sulfides from DSDP/IODP sites 134,374,376 and 654 display a wide range of variation, with values δ 34 S between -40 to +13‰ (Figure 4. 9) . Non-evaporitic marls and mudstone interbeds display the lowest δ 34 S values. Site 654 mars have on average the highest iron sulfide concentrations (on average 5705±6280 ppm) and lowest δ 34 S (average of -33±6‰). Non-evaporitic interbeds from other locations display lower iron sulfide concentration (on average 1830±1380 ppm), higher δ 34 S (on average -24±8‰), with site

Discussion

Insights from the petrographic study of sulfide bearing phases in deep basins

In marine environments, hydrogen sulfide necessary to form solid sulfide is only produced by microbial sulfate reduction; thus, the presence of chemically extractable solid sulfides in all the investigated samples confirm that sulfur in the basin was not only buried as gypsum in its oxidized form, but that a fraction of it was also microbially cycled and buried as sulfide.

Sedimentary sulfides mainly include: (i) iron monosulfides (FeS) forming through the scavenging of reactive iron by hydrogen sulfide (Eq.1) and (ii) pyrite (FeS2) resulting from the reaction of FeS with polysulfide or hydrogen sulfide (Berner, 1984;[START_REF] Rickard | Sulfidic sediments and sedimentary rocks[END_REF] (Eq.2 or 3).

Fe 2+ + H2S → FeS +2H + [START_REF] Rickard | Chemistry of Iron Sulfides[END_REF])

FeS + H2S → FeS2 +H2 (H2S pathway) [START_REF] Rickard | Kinetics of pyrite formation by the H2S oxidation of iron (II) monosulfide in aqueous solutions between 25 and 125°C: The rate equation[END_REF] (

FeS + Sx 2-→FeS2 + Sx-1 2-(Polysulfide pathway) [START_REF] Rickard | Chemistry of Iron Sulfides[END_REF]

The reduced sulfur pool used for multiple sulfur isotope analyses in this study was chemically extracted using a Chromium Reducible Solution that does not allow for a straightforward discrimination between iron monosulfide, pyrite and elemental sulfur (Fossing and Jørgensen, 1989). However, petrographic observations of our samples reveal that the large majority of the solid sulfides in our samples, regardless of the lithofacies considered, are conspicuously different from the large framboidal or euhedral pyrite crystals typically reported in [START_REF] Rickard | Sedimentary pyrite framboid size-frequency distributions: A metaanalysis[END_REF]. In addition, they consist of globules that form thin layers and seem to have encrusted the surface of other minerals (Figure 4. 7Figure 4. 8).

These features make them strikingly similar to iron monosulfide films observed in natural environments (Kotopoulou et al., 2022) or produced in laboratory experiment with microbial sulfate reducers (Berg et al., 2020;Duverger et al., 2020;Picard et al., 2018;[START_REF] Rickard | Kinetics of pyrite formation by the H2S oxidation of iron (II) monosulfide in aqueous solutions between 25 and 125°C: The rate equation[END_REF].

This is supported by the occurrence of iron oxides in addition to iron sulfides observed in our samples (Figure 4. 7) and inferred by the study of Kuroda et al. (2016) for sites 374 and 376.

Indeed, iron monosulfides instead of pyrite are generally found in environments where iron oxides available for reaction with hydrogen sulfide are in excess (Kao et al., 2004;Kotopoulou et al., 2022;Kraal et al., 2013;[START_REF] Riedinger | Sulfur Cycling in an Iron Oxide-Dominated, Dynamic Marine Depositional System: The Argentine Continental Margin[END_REF].

Solid iron monosulfides are not stable in marine environments, they get extremely rapidly oxidized in oxic waters at marine pH (Burton et al., 2009;Choppala et al., 2017) and dissolve fast to form pyrite in sulfidic environments [START_REF] Rickard | Kinetics of pyrite formation by the H2S oxidation of iron (II) monosulfide in aqueous solutions between 25 and 125°C: The rate equation[END_REF] such as encountered in classical marine sediments during early diagenesis (Goldhaber and Kaplan, 1974).

The occurrence of iron monosulfide rather than pyrite in sedimentary rocks is often interpreted to reflect low salinity environments because only a low amount of sulfate would be available to produce the reduced sulfur compound subsequently incorporated in iron sulfides (Chen et al., 2021;[START_REF] Liu | Sulfate sources constrained by sulfur and oxygen isotopic compositions in the upper reaches of the Xijiang River, China[END_REF]. This could hold true for the marl interbeds and would be supported by the low-salinity fauna they host [START_REF] Benson | The paleoecology of the ostracods of DSDP Leg 42A[END_REF][START_REF] Cita | Messinian Paleoenvironments: New Evidence from the Tyrrhenian Sea (ODP LEG 107)[END_REF]De Deckker et al., 1988;[START_REF] Popescu | Lago Mare episodes around the Messinian-Zanclean boundary in the deep southwestern Mediterranean[END_REF]Schrader and Gersonde, 1978). However, iron sulfides preserved in gypsum cannot be explained by the lack of dissolved sulfate, as the Secondly, iron monosulfides are often reported in systems with anoxic to euxinic water columns (Berner et al., 1979;Kotopoulou et al., 2022;Lyons, 1997), but this type of environment is unlikely for most of our samples, except for marls at site 654, because iron sulfide concentrations are generally very low in the samples investigated, whereas excess iron oxides seem to have been available.

Importantly, regardless of the exact nature of the iron sulfide in our samples (pyrite or iron monosulfide) the nanometric size and shape of the globules indicate early burial without further pyritization (Goldhaber and Kaplan, 1974;Hurtgen et al., 1999).

We therefore interpret iron sulfides within evaporitic deposits to be the result of early sulfidization of iron oxides where: (i) hydrogen sulfide production is sufficiently low compared to iron oxides availability, (ii) hydrogen sulfide is scavenged by reactive iron within the first few centimeters of sediments where MSR is active, and (iii) burial is sufficiently fast to prevent subsequent diagenetic formation of larger (micrometric) framboidal pyrite. This is observed in environments where all the produced hydrogen sulfide is rapidly scavenged by excess reactive iron species and allows for a thin subsurface layer of iron sulfide to form and to be rapidly buried and preserved (Hurtgen et al., 1999;Kraal et al., 2013;[START_REF] Riedinger | Sulfur Cycling in an Iron Oxide-Dominated, Dynamic Marine Depositional System: The Argentine Continental Margin[END_REF].

Such environmental conditions typically correspond to shallow water systems such as continental margins, deltas, alluvial fans, estuaries or wetlands characterized by highly dynamic depositional settings. In these environments, early diagenetic iron monosulfides instead of pyrite are commonly observed because they are formed at the surface of the sediment (<1mbsf), close to the sediment water interface and rapidly removed from the MSR Chapter 4-Late Messinian Mediterranean base level fall ___________________________________________________________________________________________ 188 zone where H2S is produced through rapid burial (Berner et al., 1979;Burton et al., 2011;Chen et al., 2021;[START_REF] Liu | Sulfate sources constrained by sulfur and oxygen isotopic compositions in the upper reaches of the Xijiang River, China[END_REF][START_REF] Riedinger | Sulfur Cycling in an Iron Oxide-Dominated, Dynamic Marine Depositional System: The Argentine Continental Margin[END_REF]Volvoikar et al., 2020).

Relationship between sedimentary settings and iron sulfide isotopic composition

The very early diagenetic character of iron sulfides entrapped in the samples studied here strongly support the pristine origin of their isotopic signature. As already proposed for micrometric individual pyrite crystals (Marin-Carbonne et al., 2022), we suggest that the δ 34 S isotopic composition measured in this study reflect the influence of redox processes involved in the micro-environment of formation of iron sulfides and not post depositional processes.

In order to interpret the iron sulfide isotopic data measured in this study, it is important to understand what controls iron sulfide δ 34 S isotopic composition in sedimentary settings.

Organoclastic and methane-based MSR zonation within marine sediments controls on iron sulfide isotopic composition

Microbial sulfate reducers preferentially metabolize the light 32 S-sulfur isotope, leading to the production of 34 S-depleted hydrogen sulfide. When the only processes involved in iron sulfide genesis is the reaction of iron with hydrogen sulfide produced by MSR [START_REF] Fike | Rethinking the ancient sulfur cycle[END_REF]Pasquier et al., 2017), its isotopic composition is shifted from that of coeval sulfate by a fractionation εMSR = -3 to -70‰ (Habicht and Canfield, 2001;Leavitt et al., 2013). The amplitude of the overall isotope fractionation εMSR is a function of the rate of MSR, itself depending on many factors, but mainly on the availability and type of reactive organic matter in environments where sulfate is not limiting (Sim et al., 2011a;Sim et al., 2011b;[START_REF] Wing | Intracellular metabolite levels shape sulfur isotope fractionation during microbial sulfate respiration[END_REF]. With increasing depth within the zone of the sediments where organoclastic-MSR (that uses organic carbon as an electron donor) is active, the sulfur isotopic composition of sulfide progress towards higher values following a Rayleigh type distillation as the sulfate pool gets progressively consumed by MSR (see the evolution of the solid lines in Figure 4. 10 as fSO4

decrease from 1 to 0). Deeper in the sediment, when porewater sulfate has been completely consumed by MSR (when fSO4=0 and δ 34 SFeS final =δ 34 SSO4 initial in Figure 4. 10), organic matter oxidation shifts from MSR to methanogenesis (Jørgensen and Kasten, 2006). In addition, in marine sediments, MSR driven by anaerobic oxidation of methane (AOM) occurs thanks to the upward diffusion of methane to the sediments depth where dissolved sulfate is present (Boetius et al., 2000;Treude et al., 2005). Hydrogen sulfide with exceptionally high δ 34 S values is often produced close to the interface where sulfate and methane are available for AOMdriven MSR: the sulfate-methane transition zone (SMTZ) (Deusner et al., 2014;Liu et al., 2021;[START_REF] Liu | Multiple sulfur isotopes discriminate organoclastic and methane-based sulfate reduction by sub-seafloor pyrite formation[END_REF]Pellerin et al., 2018). However, in sediments at water depths exceeding 10m

(and not affected by the upwards advection of sedimentary fluids), the SMTZ is generally located several meters below the sediment-water interface (Egger et al., 2018). Therefore, in most marine sediments, iron sulfides found close to the sediment water interface (<1mbsf) are unlikely to have acquired 34 S-enriched isotopic compositions (δ 34 S>-10‰), because porewater sulfate in these surficial sediments is rarely completely consumed by organoclastic MSR, and methane rarely diffuse up to such shallow sediments [START_REF] Bowles | Global rates of marine sulfate reduction and implications for sub-sea-floor metabolic activities[END_REF]Egger et al., 2018). Early diagenetic iron sulfides generally form in surficial sediments (<1mbsf) and are often depleted in 34 S by -40 to -70‰ compared to coeval sulfate (Habicht and Canfield, 2001;[START_REF] Johnston | Sulphur isotopes and the search for life: strategies for identifying sulphur metabolisms in the rock record and beyond[END_REF]Liu et al., 2020a) because: (i) the rate of MSR in most marine sediment is generally relatively low and results in large εMSR [START_REF] Bowles | Global rates of marine sulfate reduction and implications for sub-sea-floor metabolic activities[END_REF]Jørgensen et al., 2019),

(ii) the porewater sulfate concentration is high in surface sediment [START_REF] Bowles | Global rates of marine sulfate reduction and implications for sub-sea-floor metabolic activities[END_REF] because it is largely exchanging with the water column that is an effectively an infinite reservoir of sulfate (in Figure 4. 10, fSO4 stay close to 1) and further prevents (iii) the switch from MSR to methanogenesis (Egger et al., 2018;[START_REF] Riedinger | Sulfur Cycling in an Iron Oxide-Dominated, Dynamic Marine Depositional System: The Argentine Continental Margin[END_REF]. These features are particularly prevalent in surficial sediments of basins exceeding 200 m depth [START_REF] Bowles | Global rates of marine sulfate reduction and implications for sub-sea-floor metabolic activities[END_REF]Egger et al., 2018). For these reasons, and because they didn't undergo extensive In comparison to the picture exposed above, the δ 34 S of iron sulfides measured in this study displays heavier values (-13±16‰) and large amplitude variations (>50‰). Although some of the iron monosulfides display a large offset (shifted by a maximum of -61‰) from coeval Messinian Sea water sulfate (δ 34 SMESS≈22‰) (Masterson et al., 2016), more than half of the values measured here have δ 34 S> -5‰. These values are not expected for typical early diagenetic iron sulfide if only MSR expressing large fractionation was at play. Hypersaline environments as prevailing during stage 3 calcium sulfate minerals precipitation have been suggested to favor the expression of reduced fractionation factors during MSR (εMSR between -9 and -20‰;

( Avrahamov et al., 2014)). Although this could explain part of the most enriched values from our dataset, overall, the measured iron sulfide from calcium sulfate deposits are offset from coeval sulfate by εnet= -8 to -55‰, covering almost the entire range of possible fractionation expressed during MSR, and not only small values. Therefore, additional processes seem to be needed to explain the large amplitude variations displayed by the isotopic composition of iron sulfide extracted from both gypsum and marls interbeds.

Water column depth control on sedimentary settings and isotopic composition of iron sulfides

In a recent study, Pasquier et al. (2021a) proposed that the thickness of the water column (X-Y , in our nomenclature of Figure 4. 2)is a major control on the isotopic composition of preserved iron sulfide. In the eastern New Zealand sedimentary system, they showed that large amplitude variations in pyrite δ 34 S ∼70‰ in sediments of the shelf (water depth of In this work, we build on the observation made by Pasquier et al., 2021a in the New Zealand sedimentary system. Because the production of enriched iron sulfide in 34 S and their distance to the sediment water interface is mainly controlled by the sedimentary regime of the basin considered, we suggest that in most deep water marine environments, high variability in δ 34 S isotopic values of iron sulfides couldn't be observed. Indeed, the production or vertical transport of 34 S-enriched sulfide to surficial sediments (<1mbsf) occur through 4 main controls:

• the significative consumption of surficial porewater sulfate (fSO4<1 in Figure 4. 10) due to (1) high sedimentation rates allowing for the efficient disconnection of surficial porewater sulfate from the overlying water column (Pasquier et al., 2017) or (2) high organic carbon fluxes or availability that lead to high MSR rates : in that case sulfate consumption in the porewater is faster than its replenishment by diffusion from the water column (Pasquier et al., 2021b) • the mixing of deep 34 S-rich iron sulfide with surficial sediments through (3) large amplitude and frequent sediment reworking and resuspension (Aller et al., 2010;Pasquier et al., 2021a) Chapter 4-Late Messinian Mediterranean base level fall ___________________________________________________________________________________________ 193

• the production of 34 S enriched iron sulfides caused by (2) high organic carbon fluxes or availability that leads to higher MSR rates and causes smaller fractionation between sulfide and sulfate (low εMSR in Figure 4. 10), (Mengdi et al., 2022) or (4) the occurrence of reoxidative reactions that can further fractionate sulfur isotopes [START_REF] Findlay | Sulfide oxidation affects the preservation of sulfur isotope signals[END_REF]Habicht and Canfield, 2001;[START_REF] Habicht | Sulfur isotope fractionation during bacterial reduction and disproportionation of thiosulfate and sulfite[END_REF]Pellerin et al., 2015;Pellerin et al., 2019).

As a matter of fact, (1) high sedimentation rates, (2) high organic carbon fluxes or availability

(3) large and frequent episodes of mixing or resuspension or (4) the occurrence of an important reoxidative sulfur cycling are all features that are unlikely to occur in deep marine basins and typically found in marginal and shallow water basins rather than in deep basins (Aller, 2014;Aller et al., 2010;[START_REF] Jørgensen | Sulfur biogeochemical cycle of marine sediments[END_REF]Middelburg et al., 1997;[START_REF] Restreppo | Oceanic sediment accumulation rates predicted via machine learning algorithm: towards sediment characterization on a global scale[END_REF].

Thus, deep water marine basins are unlikely to be highly dynamic sedimentary systems, and couldn't host a biogeochemical sulfur cycle favorable to the formation of iron sulfide displaying high amplitude variations and on average high δ 34 S. Conversely, the reduced water column thickness and the lateral proximity of shallow water basins to the land favor these features.

However, the water depth at which sedimentation rate, organic matter fluxes, sediment mixing or sulfide reoxidation are negligible effects and stop to shape the δ 34 S of surficial sediment iron sulfides is poorly constrained. 4.5.2.4 Constrains on the water depth of deposition of iron sulfide within stage 3 deposits at site 134, 374,376 and 654 Considering that stage 3 deposits studied in this work were retrieved at 1965 m to 4560m mbsl, and considering the empirical depth limit of 1500mbsl determined above, we thus suggest that the iron sulfides we measured formed under a water column (X'Y') that was thinner than that at which they've been sampled (X-Y). Based on our compilation (Figure 4. [START_REF] Brüchert | Biogeochemical and Physical Control on Shelf Anoxia and Water Column Hydrogen Sulphide in the Benguela Coastal Upwelling System off Namibia[END_REF], we propose that the iron sulfide we measured in this study, and for which half of the values exceed -5‰ couldn't have formed at deeper depth than 1500m. Although it seems difficult to make of this assumption a rule for every sedimentary context, we detail below why we think this general rule work in the very specific case of the MSC stage 3 deposits we studied here. 1) sedimentation rates in deep Mediterranean basins are one to two orders of magnitude lower (0.03 to 0.12m/kyrs) (Basso et al., 2004;Garcia-Orellana et al., 2009;van Santvoort et al., 1996;Zúñiga et al., 2007) than what has been observed to generate a rapid and efficient disconnection of porewater sulfate from the overlying water column (0.5 to 2.5 m/kyrs)-sufficient to increase the δ 34 S of the sulfide accumulating in this micro environment (Liu et al., 2021;Pasquier et al., 2017;Volvoikar et al., 2020).

Further, none of the sites studied here is located close to the mouth of a large river or in particular sedimentary settings (see chapter 2 for detail context of each well) that could justify exceptionnally high sedimentation rates (Volvoikar et al., 2020) at any of the sites if they were under 2000 to 4000m of water column. respectively. The fact that MSR rates are necessarily low in deep basins seems to hold true for deep stratified restricted basins, where the amount of organic carbon reaching the seafloor is high, but where organic matter is possibly poorly available for microbial sulfate reducers. Indeed, the iron sulfide formed in deep stratified basins preserving large amounts of organic matter have often constant and 34 S-depleted isotopic compositions, as observed in Pleistocene deep Mediterranean basin sapropels [START_REF] Passier | Modes of sapropel formation in the eastern Mediterranean: some constraints based on pyrite properties[END_REF][START_REF] Scheiderich | Molybdenum isotope, multiple sulfur isotope, and redox-sensitive element behavior in early Pleistocene Mediterranean sapropels[END_REF], and especially when sulfate concentration is high [START_REF] Gomes | Sulfur isotope fractionation in modern euxinic systems: implications for paleoenvironmental reconstructions of paired sulfatesulfide isotope records[END_REF]. In addition, even though particularly high rates of MSR have already been observed at the bottom of deep water hypersaline and anoxic marine basins (Joye et al., 2009), isotopic measurement of reduced sulfur species produced in the Mediterranean deep brine pools display lower δ 34 S values than -15‰ (Henneke et al., 1996).

Further, high TOC content do not seem to be supported for most of the samples studied here (McIver and RD, 1973;Pierre and Rouchy, 1990;Sigl et al.), especially for the calcium sulfate minerals that have a whitish aspect. This means that if high organic content was driving the enrichment in heavy isotopes, given that marls are richer in Chapter 4-Late Messinian Mediterranean base level fall ___________________________________________________________________________________________ 198 organic matter than gypsum (see section 4.2), it is expected that marls would bear iron sulfides richer in 34 S than gypsum. The opposite trend, however, is observed in our dataset (Figure 4. 9). Another way to produce 34 S-enriched iron sulfide in deep marine settings would be that they were formed close to the SMTZ as observed for the most 34 S-enriched iron sulfide of our compilation (see Figure 4. 11). Since the iron sulfide we extracted are contained in gypsum, the sedimentary column was oversaturated with respect to sulfate at the time of iron sulfide formation, precluding for the existence of a shallow SMTZ at the sites we studied. Thus, low organic content and the weakly positive to weakly negative δ 13 C of the carbonates phases associated to stage 3 calcium sulfate does not support the occurrence of methane seepage at any of these sites during MSC stage 3 [START_REF] Fontes | Oxygen, Carbon, Sulfur and Hydrogen stable isotopes in carbonate and sulfate mineral phases of Neogene evaporites, sediments and in interstitial waters[END_REF][START_REF] Ricchiuto | Stable Isotopic investigation of Messinian sulfate samples from DSDP. LEG 42. Eastern Mediterranean Sea[END_REF]Weinheimer, 1978).

3) Substantial enrichment of sedimentary sulfide isotopic composition by mixing is observed when episodic resuspension of sediment at a larger depth, in highly dynamic coastal sedimentary environment, brings together 34 S-enriched sulfide pool with newly formed 34 S-depleted sulfide [START_REF] Fike | Rethinking the ancient sulfur cycle[END_REF]Pasquier et al., 2021a). The mixing of sediments in the deep Mediterranean is however controlled by bioturbation that has a relatively homogenous and constant effect in time and space on the mixing of sediment and affect only the first 10cm of sediment (Barsanti et al., 2011). Therefore, large amplitude sediment resuspension seems unlikely to have driven iron sulfide isotopic composition in the deep basins of the Mediterranean we consider in this study, especially since these basins are separated by several hundreds of kms. 4) Marine sediments in basins ranging from 400m to 2200m have been shown not to be affected by substantial reoxidation of sulfide [START_REF] Turchyn | Oxygen isotopic composition of sulfate in deep sea pore fluid: evidence for rapid sulfur cycling[END_REF]. Indeed, reoxidation of sulfide is mainly driven by two factors : (i) the availability of large Chapter 4-Late Messinian Mediterranean base level fall ___________________________________________________________________________________________ 199 amount of oxidants, (iron, manganese..) either brought by large riverine inputs or atmospheric deposition, that mainly depends on the proximity to their sources, i.e. the land (Blonder et al., 2017;[START_REF] Boyko | Iron and sulfur speciation and cycling in the sediments of marine systems located in arid environments: the northern Red Sea[END_REF][START_REF] Mills | Geochemical evidence for cryptic sulfur cycling in salt marsh sediments[END_REF] or (ii) the occurrence of microbial sulfide oxidizers that flourish in the presence of free hydrogen sulfide in highly productive zones of the ocean (mainly focused on coastal areas (van Vliet et al., 2021;Wasmund et al., 2017)) [START_REF] Jørgensen | Sulfur biogeochemical cycle of marine sediments[END_REF] or hydrothermal vents (Kotopoulou et al., 2022).

From this theoretical understanding of the sedimentary dynamics governing the deep basins of the Mediterranean, it seems unlikely that the gypsum deposits studied here formed in basins exceeding 2000m depth, which is in contrast with their current depth.

Accordingly, we suggest that the petrographic features (Figure 4. 7Figure 4. 8) and the large isotopic variability displayed by the iron monosulfides preserved in our samples (Figure 4. 9Figure 4. 11), were shaped by high amplitude oscillations of sedimentary settings due to a dynamic sedimentary regime, characteristic of water level oscillations in relatively shallow water depths basins (<1500mbsl) (Middelburg et al., 1997).

So far, evidence for a shallow water level in the deepest part of the Mediterranean during stage 3 comes from erosive features deduced from the study of seismic makers (Lofi et al., 2011a;[START_REF] Madof | Discovery of vast fluvial deposits provides evidence for drawdown during the late Miocene Messinian salinity crisis[END_REF]Micallef et al., 2018a) or from the rare occurrence of shallow water fauna in marl interbeds [START_REF] Popescu | Lago Mare episodes around the Messinian-Zanclean boundary in the deep southwestern Mediterranean[END_REF]. Our study provides an independent, additional support arguing against continuously high-water levels during stage 3 deposition. It supports relatively low water levels (<1500m) during stage 3, not only during erosive phases as evidenced in seisimic studies, but also during deposition of stage 3 sedimentary layers. This also disproves one of the high-water levels scenario for stage 3. In the "deep basin deep water" model, dense brine cascading and dissolution of evaporites were supposed to explain [START_REF] Gvirtzman | Intra-Messinian truncation surface in the Levant Basin explained by subaqueous dissolution[END_REF]Roveri et al., 2014c). Such a phenomenon in basins exceeding 2000m depth, without at least intermittent water level fall, would not be able to generate the iron sulfide isotopic data we measure at any of the sites considered in this study.

Implications for isostasy based paleo-depths reconstructions of stage 3 layers

The restricted paleo-water depths at the sites studied here towards the end of the MSC have large implications regarding the scenario ruling the MSG formation. This raises questions about the hydrological conditions necessary to match the maximal 1500m water thicknesses (X'Y'max) constrained by our isotopic data: Did shallow water levels could have take place without a significant sea level draw down i.e., could the bathymetry during stage 3 have been shallow enough to explain our results without a drawdown? And if a drawdown was necessary, what was its amplitude? Was it homogenous in all the Mediterranean given that the Sicily Sill separating its East and West basins is relatively shallow? Using X'Y'max the maximum water column thickness estimate of 1500m deduced from our isotopic data, we tried to estimate the bathymetry Y' of each of the basins considered in this study in order to understand the conditions necessary to induce shallow water (<1500m) deposition at the investigated sites. To achieve this goal, the magnitude of the various components of vertical motion have been estimated following the method described in section 4.3.3. Messinian evaporites records open marine conditions with high productivity, and records sediment instability due to tilting and faulting of the crustal block on which the well was drilled. This means that before and after the MSC, water levels at site 654 were high; thus, the shallow water level inferred by the isotopic record at this site should have been acquired during the MSC. Due to the position of the well on a margin undergoing recent extension (up to 6 Ma) with a reported heat flow (50mW/m 2 ) which is in accordance with a stretching factor of 1.4-1.5, the original depth of site 654 stage 3 deposits at the end of the Miocene is hard to constrain, but is unlikely to have been shallower than 1500 m if no major sea level drop affected this site. A depth of deposition shallower than 1500 m would have required post- Messinian tectonic subsidence in the order of 700-800 m, which is unrealistically high for nonoceanic crust with an extension factor in the reported range (estimated at <100 m for continental stretching model by [START_REF] Ricchiuto | Stable Isotopic investigation of Messinian sulfate samples from DSDP. LEG 42. Eastern Mediterranean Sea[END_REF].

Paleodepth reconstructions

The Ionian basin, where site 374 samples were recovered, is formed of old oceanic crust not likely to have undergone significant alteration by tectonics or thermal cooling since the Miocene, as Pliocene-Quaternary deposits at the site show a mostly uniform, parallel character [START_REF] Camerlenghi | Seismic markers of the Messinian salinity crisis in the deep Ionian Basin[END_REF]. The site was targeted for its position in the central part of the Messina abyssal plain and thick Pliocene-Quaternary sedimentary sequence, but it cannot be excluded that the vicinity of the Calabrian and Mediterranean ridge subduction systems affected the depth of the Ionian abyssal plain. Analysis of benthic formanifera assemblages showed consistent Pliocene-Quaternary upper Mesobathyal (>1500 m minimum depth), including some zones containing Middle Mesobathyal (>1800 m depth) assemblages (Wright, 1978). This means that directly after the MSC, the water depth at site 374 exceeded 1500 to 1800m. Compaction of the pre-MSC sediments by Plio-Quaternary sedimentation can be assumed to be small as the thick evaporite layer is most likely to already have caused the bulk of compaction before the onset of the Pliocene. This leaves only a minor (~180 m) isostatic subsidence component altering the depth of the Miocene-Pliocene boundary. A very significant depth of up to 4000m can be assumed at the end of the MSC if no significant base level fall occurred at the acme of the crisis, barring large-scale post-Messinian foundering of the basin (of >2km), with this subsidence not being reflected in changes in the fossil assemblage.

The Florence Rise, where ODP wells 376 (and complementary well 375 some kilometers to the south) was drilled is the Westward continuation of the Cyprus arc, a topographic high bounded by south-verging thrusts marking the boundary between the Africa The region has undergone transpressional reactivation related to the westward escape of Anatolia since the Pliocene ( Mart and Ryan, 2002). Erosion has been suggested to have occurred during the MSC as attested by the very thin salt units in this basin compared to the Antalya Basin to the north and Herodotus abyssal plain to the south (Sellier et al 2013). We can confidently assume that the Florence Rise has been uplifted since the MSC, marking its current depth of the Miocene-Pliocene boundary at 2240 mbsl as a lower depth limit for the end Miocene, even though uplift of the Messinian units at site 376 was likely smaller that on Cyprus, where deformed Messinian evaporites are found onshore. Isostatic motions due to post-Messinian sedimentation would have played a minor role in this region compared to tectonic deformation on the ridge.

The depth range estimate of stage 3 deposits for the "non-drawdown" scenario (>1500m to >4000m) are way over the water depths limits imposed by the extreme amplitude variability of our iron sulfide isotopic data, that preclude the deposition of the stage 3 deposits studied here at greater water depths than 1500m maximum.

Scenario 2: drawdown

To match the water column thickness X'Y'max implied by our isotopic measurements with isostasy-based depth estimates, a draw down, that would have allowed for its reduction through both water removal and isostatic rebound, is necessary (Figure 4. 5). We can now examine which scenario of base level fall during the last stage of the MSC fit best with our observations for each site. If we consider that the present day Sicily Sill -that separates the western from the eastern basins -was not at a drastically different position than nowadays, that is of ∼430m (Blanc, 2006) it means that any large amplitude Mediterranean sea level drop Chapter 4-Late Messinian Mediterranean base level fall ___________________________________________________________________________________________ 204 would have allowed the disconnection of eastern and western basins [START_REF] Andreetto | Freshening of the Mediterranean Salt Giant: controversies and certainties around the terminal (Upper Gypsum and Lago-Mare) phases of the Messinian salinity crisis[END_REF][START_REF] Lofi | Erosional processes and paleo-environmental changes in the Western Gulf of Lions (SW France) during the Messinian Salinity Crisis[END_REF]Micallef et al., 2018a). We therefore considered depth estimates separately for sites 134 and 654 in the western Mediterranean and for sites 374 and 376 in the eastern Mediterranean.

Western Mediterranean

Based on paleobathymetric reconstruction of Heida et al. ( 2021) and their estimations for a sea level (XX') drawdown of -1100m of amplitude, the water depth at site 134 would have been of ∼1100m to 1200m if a drawdown operated. Following the same scenario of 1100m base level fall (Dwn) for site 654, the paleobathymetry Y' should have been at about 1200 m before compensating for isostatic rebound, implying a water column a few hundred meters deep during part of stage 3. The water column thicknesses estimate for site 134 (X'-Y') is close to the threshold water depths possibly enabling the production of excessively enriched 34 S iron sulfide. Thus, our result further supports that a base level drop of -1100m in the Western Mediterranean is a minimum estimate in order to fit the evidences observed as previously stated by numerous studies on seismic markers of the western Mediterranean [START_REF] Lofi | Erosional processes and paleo-environmental changes in the Western Gulf of Lions (SW France) during the Messinian Salinity Crisis[END_REF]Lofi et al., 2011a;Raad et al., 2021), For site 654 in the Tyrrhenian basin, the paleodepth estimates are uncertain due to tectonics. However, assuming our estimate in section 4.3.3.1 of a non-drawdown paleodepth of at least 1500-2000 m is correct, a water level drop of around 1100 m would have allowed the basin to have a water depth <500m, that would be consistent with the dynamic sedimentary settings inferred by the highly variable isotopic values found at sites 654.

.2.2 Eastern Mediterranean

In the Ionian Basin at site 374, the no-drawdown scenario yields a bathymetry (Y') similar to its current depth (Y) of 4000 m, not accounting for possible foundering of the basin due to the vicinity of the Calabrian and Mediterranean ridge system [START_REF] Camerlenghi | Uncovering the Mediterranean Salt Giant (MEDSALT)-Scientific Networking as Incubator of Cross-disciplinary Research in Earth Sciences[END_REF].

To match the maximum water thickness estimate (X'Y'max) suggested by our isotopic observations, the drawdown would require a water depth reduction of 2500 m (X'Y'full -X'Y'max in Figure 4.4Erreur ! Source du renvoi introuvable.), corresponding to a water level fall XX' of around -1700 m. This is not far from the result by Micallef et al. (2018b) of a more detailed isostatic restoration of paleoshorelines and shore platforms on the Malta Escarpment on the western limit of the Ionian Basin, where they infer a shoreline level 1800-2000 m below modern sea level.

However, a drawdown of such magnitude would imply that site 376 was possibly subaerially exposed, accounting a non-drawdown depth of around 2300 m. This would leave a water column of only a few hundred meters, possibly explaining the proposed erosion of the Florence Rise at this time. A minimum water level reduction (X'Y'full -X'Y'max in Figure 4.4) at site 376 would be approximately of 2300-1500 = 800 m, corresponding to a water level drop (XX') of around 550 m, to reach our reference 1500 m water depth. However, the isotopic composition of iron sulfides within calcium sulfate deposits measured at site 374 cannot be explained by such a limited sea level fall (XX').

The discrepancy in required water level drop estimation at the two Eastern Mediterranean sites can be explained by: -a deeper bathymetry at site 376, exceeding our 2300 m minimum estimate.

-a shallower bathymetry at site 374, still exceeding 1800 m but less than its current 4000 m. An unidentified source of vertical motion, possibly related to geodynamic structural highs, at present-day depths of 4.4 km below sea level (Camerlenghi et al. 2019).

-differences in water level during lowstand between the Messina abyssal plain and the Florence Rise controlled by a hydrological sill shallower than the Florence Rise at, for example, the Mediterranean Ridge (as proposed by [START_REF] Camerlenghi | Seismic markers of the Messinian salinity crisis in the deep Ionian Basin[END_REF], separating the Eastern Medierranean into two separate subbasins during lowstand.

-Variations in water level during the lowstand temporarily exposing the Florence Rise to erosion between phases of shallow water deposition, and phases with a water level above and below the 1500 m reference on the Messina abyssal plain.

These conclusions have to be weighed considering the fact that the true magnitude of tectonic vertical motion at site 376 is not well constrained. Site 374 is formed of old oceanic crust and vertical motions since the MSC are not obvious from its seismic stratigraphy or fossil assemblage, but we cannot exclude that it underwent significant alteration by geodynamic processes since the Miocene. We suggest a plausible water height reduction of at least -2500 m (X'Y'full -X'Y'max in Figure 4.4) corresponding to a water level drop XX' of at least -1700 m in the eastern Mediterranean before the deposition of the uppermost part of stage3 deposits.

Site 376 was most probably particularly shallow, in agreement with the fact that most of the calcium sulfate deposits at this site consist of clastic deposits.

The hypothesis of a large base level fall having affected the eastern Mediterranean is further supported by seismic markers in the Ionian basin (where site 374 is located), suggesting its disconnection from other basins during an important draw down [START_REF] Camerlenghi | Uncovering the Mediterranean Salt Giant (MEDSALT)-Scientific Networking as Incubator of Cross-disciplinary Research in Earth Sciences[END_REF] Chapter 4-Late Messinian Mediterranean base level fall ___________________________________________________________________________________________ 207 Micallef et al., 2018a). Further, large fluvial deposits offshore southern Cyprus area [START_REF] Madof | Discovery of vast fluvial deposits provides evidence for drawdown during the late Miocene Messinian salinity crisis[END_REF] and the occurrence of paleosols and erosional features onshore Cyprus [START_REF] Orszag-Sperber | The onset of the Messinian salinity crisis: insights from Cyprus sections[END_REF][START_REF] Rouchy | Paleoenvironmental changes at the Messinian-Pliocene boundary in the eastern Mediterranean (southern Cyprus basins): significance of the Messinian Lago-Mare[END_REF]) (site 376 is located offshore western Cyprus) suggest that the entire area had extremely shallow waters or even was periodically exposed during stage 3.

Conclusion

In this work, we show that the isotopic composition of iron sulfides can be useful to decipher the thickness of the water column under which they form. While, to our knowledge, there are no direct water level proxies during stage 3 gypsum deposition, the results of the present study provides a strong evidence for shallow water levels (<1500mbsl) in the abyss of the Mediterranean during the formation of gypsum. In this study we demonstrate that stage 3 deposits across a West East transect of the Mediterranean: in the Balearic Abyssal plain (site 134), in the Upper Sardinia margin (site 654), in the Ionian Abyssal Plain (site 374) and in the Antalya basin (site 376), were all deposited under a water column shallower than 1500m.

Comparing these estimates with the current water depth of these sites (between 2000 and 4000m depth) and performing isostasy-based paleo depth reconstructions, revealed that a significant drawdown is required to have occurred to explain the shallow water environments inferred by our isotopic measurements. These findings are completely independent of the seismic record and further support existing models for an important water level fall of the Mediterranean, with at least 1.1km of drawdown of the Western Mediterranean basin water level and approximately 2.5km in the Eastern Mediterranean. These estimates are subject to significative uncertainty as stated in the sections above, but support the disconnection Indeed, the timing of such large amplitude base level fall of the Mediterranean is uncertain.

From our method we cannot precisely determine at what period this shallowing started and stopped and whether or not it was a unique event or several episodes of water level oscillation. Nevertheless, the present findings agree with the repeated high amplitude water level fluctuations of the Mediterranean during stage 3, as proposed in our recent, work in the intermediate depth basin of Eraclea Minoa in Sicily (see Appendix E for full text, (Andreetto et al., 2022a).

Introduction

The variations in isotopic values of sulfur bearing sedimentary archives (i.e., δ 34 SSO4 of gypsum, barite or carbonate associated sulfate and δ 34 SFeSx of iron sulfide or pyrite), and the net fractionation factor 34 εnet ≈ δ 34 SFeSxδ 34 SSWSO4 between co-eval seawater sulfate and sulfide bearing minerals have often been used to deduce the long-term shifts in fluxes of oxidized (sulfate) and reduced (sulfide) sulfur species to and from the ocean (e.g. Canfield, 2004;Halevy et al., 2012;[START_REF] Wortmann | Rapid variability of seawater chemistry over the past 130 million years[END_REF]Yao et al., 2018). This is important because changes in the relative size of these fluxes -largely driven by the burial and dissolution of sulfate and sulfide bearing minerals (Figure 5. 1) play a crucial role in regulating the long-term earth surface redox state (Berner, 1989;Ebelmen., 1845;Garrels, 1984). However, recent studies have demonstrated that the isotopic composition of sulfur minerals does not necessarily capture the biogeochemical state of the oceanic reservoir at the time of their precipitation

Chapter 5-: Unraveling the roles of sulfide oxidation and of variable fractionation factors during microbial sulfate reduction ___________________________________________________________________________________________ 210 but rather the partitioning of sulfur species within smaller reservoirs (such as the sediment porewater or semi-enclosed restricted basins) and on shorter time scales (Figure 5. 1) [START_REF] Fike | Rethinking the ancient sulfur cycle[END_REF]Houghton et al., 2022;Pasquier et al., 2021a;Pasquier et al., 2021b;Pasquier et al., 2017).

In these reservoirs, the isotopic composition of sulfur bearing minerals is affected by both the relative size of the fluxes of sulfur consumed by biogeochemical redox reactions, as well as by the magnitude of isotope fractionation expressed during these processes (Figure 5. 1). Biogeochemical cycling of S typically takes place in spatially restricted areas that can vary in dimensions from single marginal basins to porewater environments (e.g., Bryant et al., 2020;Thomazo et al., 2019). In these environments, the residence time of sulfate, and therefore the (hours/days/years?) than those that control sulfate concentration in the global ocean (in the order of Myrs (Holser and Kaplan, 1966;Lécuyer, 2016)). Therefore, the isotopic composition of sulfate from these environments can be different from that of the global ocean, and variations in sulfur isotopic composition of authigenic minerals can reflect fluctuations in sedimentary biogeochemical cycling, rather than global shifts in oceanic biogeochemistry.

Whereas it questions the use of the sulfur isotopic variation of S-bearing minerals in the characterization of global oceanic sulfur cycle (Jones and Fike, 2013), these findings make sulfur isotopes a promising proxy to finely reconstruct local paleoenvironmental settings that control the isotopic composition of sulfur minerals. Specifically, iron sulfides are good candidates to investigate the control of a large range of environmental conditions on sulfur isotopic composition, because different from sulfate bearing minerals that form in specific environmental conditions, they are ubiquitous in marine sedimentary settings.

In Chapter 4 of this manuscript, we presented the isotopic composition of iron sulfide extracted from gypsum formed in deep basins of the Mediterranean during the terminal phase of the Messinian Salinity Crisis (Cita and Ryan, 1978;Haq et al., 2020;Hsü et al., 1973;Lofi, 2018). Whereas we demonstrated that the isotopic composition of iron sulfide were shaped by local depositional parameters, fundamental questions remain to identify which where the processes involved and how environmental settings controlled the large isotopic variations they record.

Firstly, in the restricted reservoirs where the biogeochemical cycling of sulfur occurs, when a significant fraction of the sulfate pool (fSO4; Figure 5. 2) has been consumed by MSR, it generates the Rayleigh distillation of isotopes in porewaters, causing the accumulated sulfide to be enriched in the heavy isotope 34 S (Gomes and Hurtgen, 2013) (Figure 5. 2). This phenomemenon, called sulfide pooling has been often linked to an increase in the rates of sedimentation, or of the rates of microbial sulfate consumption that favour the isolation of the sulfate reservoir being consumed by MSR (Houghton et al., 2022;Pasquier et al., 2021b).

Secondly, iron sulfide with isotopic composition enriched in 34 S can also be produced if the fractionation factor during MSR, 34 εMSR ≈ δ 34 SH2Sδ 34 SSO4, is reduced (Figure 5. 2), for example when cell specific rates of MSR (csSRR) are high because organic carbon availability is important (Jiang et al., 2022;Leavitt et al., 2013;Sim et al., 2011a;Sim et al., 2011b). Whether

or not this effect occurs in natural sediments -and is preserved in the rock record -is debated [START_REF] Jørgensen | Sulfur biogeochemical cycle of marine sediments[END_REF]. This is an important question because, should this effect be relevant, it The deposits accumulated in the Mediterranean basins at the end of the MSC, constituted of calcium sulfate and marls interbeds, are ideal candidates to understand the relative importance of these two phenomena -Rayleigh distillation and/or varying fractionation factors during MSR. Indeed, iron sulfide preserved within calcium sulfate minerals were necessarily formed and preserved in an environment saturated with respect to gypsum, and thus rich in dissolved sulfate Further, the measurement of the calcium sulfate minerals isotopic composition provides an estimate of the sulfate isotopic composition that was biogeochemically cycled in order to produce the iron sulfide entrapped in them. In a sulfate saturated environment where coeval sulfate and sulfide sulfur isotopic measurement are available, 34 εnet can be calculated. Therefore, the influence of a Rayleigh distillation (where 34 εnet should be constant but both δ 34 SSO4 and δ 34 SFeS evolve towards isotopically heavier values) is theoretically easily distinguishable from a change in fractionation factor 34 εMSR during MSR (where δ 34 SSO4 would not necessarily vary but where 34 εnet and δ 34 SFeS would be expected to fluctuate) (Jones and Fike, 2013) (Figure 5. 2).

These Messinian deposits thus provide a rare opportunity to detect whether or not variations in the fractionation factor during MSR can be preserved in the geological record (Jiang et al., 2022;Leavitt et al., 2013) or if enrichment in 34 S isotope of sedimentary sulfide are necessarily due to Rayleigh distillation and sulfide pooling as proposed by most of the current studies on

Chapter 5-: Unraveling the roles of sulfide oxidation and of variable fractionation factors during microbial sulfate reduction ___________________________________________________________________________________________ 214 iron sulfide isotopic composition [START_REF] Jørgensen | Sulfur biogeochemical cycle of marine sediments[END_REF][START_REF] Masterson | Interpreting multiple sulfur isotope signals in modern anoxic sediments using a full diagenetic model (California-Mexico margin: Alfonso Basin)[END_REF]Pasquier et al., 2021b).

Nevertheless, other processes such as the mixing during diagenesis between various sulfur pools having different isotopic composition, or re-oxidative processes can overprint the isotopic signature of MSR [START_REF] Findlay | Sulfide oxidation affects the preservation of sulfur isotope signals[END_REF]Pasquier et al., 2021a), leading to 34 εnet values that cannot be ascribed exclusively to MSR ( 34 εnet ≠ 34 εMSR). Further, in some cases, re-oxidative processes can lead to the production of reactive sulfur species not preserved in the sedimentary record (Findlay and Kamyshny, 2017), and thus not measurable. Whereas this phenomenon is difficult to study in the rock record, it still has the potential to influence the fluxes and isotopic composition of the dissolved sulfur species that are ultimately preserved as minerals (Johnston et al., 2005;Pellerin et al., 2019;[START_REF] Zerkle | Fractionation of multiple sulfur isotopes during phototrophic oxidation of sulfide and elemental sulfur by a green sulfur bacterium[END_REF]Zerkle et al., 2016) (Figure 5. 1).

Together with the oxygen isotopic composition of sulfate δ 18 OSO4, the addition of 33 S minor sulfur isotopes measurement to the classical 34 S and 32 S isotopes allows to overcome this issue, to detect the occurrence of re-oxidative processes and to distinguish them from concurrent MSR [START_REF] Antler | The Sedimentary Carbon-Sulfur-Iron Interplay -A Lesson From East Anglian Salt Marsh Sediments[END_REF]Böttcher et al., 1998;Brunner et al., 2005;Farquhar et al., 2003;[START_REF] Johnston | Sulphur isotopes and the search for life: strategies for identifying sulphur metabolisms in the rock record and beyond[END_REF]Ono et al., 2006;[START_REF] Zerkle | Fractionation of multiple sulfur isotopes during phototrophic oxidation of sulfide and elemental sulfur by a green sulfur bacterium[END_REF]. In order to (i) differentiate

Biogeochemistry of the Mediterranean Sea during the MSC

Sulfate is a major constituent of gypsum and anhydrite deposits that represent a significant percentage of the evaporitic deposits formed in the Mediterranean during the late Miocene Messinian Salinity Crisis (Cita and Ryan, 1978;Haq et al., 2020;Hsü et al., 1973;Lofi, 2018).

The impact of this dramatic event on short term sulfur cycling has often been overlooked in studies measuring δ 34 S sulfur isotope composition of the sulfate minerals constituting this salt giant, using it only to argue in favor of the marine origin of the sulfate forming gypsum [START_REF] Evans | Coupled measurements of delta O-18 and delta D of hydration water and salinity of fluid inclusions in gypsum from the Messinian Yesares Member, Sorbas Basin (SE Spain)[END_REF]Lugli et al., 2007). However, large evidences exist for the modifications of the biogeochemistry of the Mediterranean basin following its hydrologic restriction from the Atlantic Ocean (Bulian et al., 2021;Dela Pierre et al., 2012;Kouwenhoven et al., 2003;[START_REF] Kouwenhoven | Paleoenvironmental evolution of the eastern Mediterranean during the Messinian: constraints from integrated microfossil data of the Pissouri Basin (Cyprus)[END_REF]Natalicchio et al., 2019;Sabino et al., 2022;[START_REF] Sabino | The response of water column and sedimentary environments to the advent of the Messinian salinity crisis: insights from an onshore deep-water section (Govone, NW Italy)[END_REF].

Other studies considering also oxygen isotopes of sulfate (δ 18 OSO4) suggest that part of the sulfate forming Messinian evaporites underwent re-equilibration with ambient water, a process that is only possible when sulfate is biogeochemically cycled [START_REF] Aloisi | The geochemical riddle of "low-salinity gypsum" deposits[END_REF]García-Veigas et al., 2018;Lu et al., 2001;[START_REF] Waldeck | Calibrating the triple oxygen isotope composition of evaporite minerals as a proxy for marine sulfate[END_REF]. Further, in the Vena del Gesso marginal basin (Northern Apennines) our previous work suggested that this biogeochemical sulfur cycling has been hidden from the δ 34 S of sulfate record because microbial sulfate reduction into sulfide was overprinted by the complete reoxidation of sulfide i.e a cryptic sulfur cycling was ongoing during gypsum formation (Guibourdenche et al., 2022)(see chapter 3 for the full text or Appendix F for the published version of the paper).

These studies almost all concern Primary Lower Gypsum deposits (PLG) formed at the onset of the MSC, mostly concentrated in small, shallow, marginal and restricted basins of the Mediterranean. Nonetheless, gypsum deposition was not restricted to small, shallow marginal

Chapter 5-: Unraveling the roles of sulfide oxidation and of variable fractionation factors during microbial sulfate reduction ___________________________________________________________________________________________ 216 basins [START_REF] Manzi | A new facies model for the Upper Gypsum of Sicily (Italy): chronological and paleoenvironmental constraits for the Messinian salinity crisis in the Mediterranean[END_REF] but also took place in the broader Mediterranean, including its deepest, central basins (Andreetto et al., 2021). Indeed, a significant part of MSC calcium sulfate deposits are lying in the offshore domain of the Mediterranean in intermediate ( 200to 1000m) to deep basins (>1000m) (Haq et al., 2020;Lofi, 2018). The thick evaporitic sequence that lie offshore on top of halite deposits, mostly in the western Mediterranean has been mainly studied by seismic imagery, and form in deep basins of the Mediterranean a seismic unit called Upper Unit (UU) (Lofi, 2018;Lofi et al., 2011;Roveri et al., 2014). The uppermost part of some of these deposits have been drilled by ODP/DSDP expeditions and are interpreted to correspond to Upper Gypsum deposits that are constituted by gypsum and marls alternation and formed during the third and last phase of the MSC [START_REF] Manzi | A new facies model for the Upper Gypsum of Sicily (Italy): chronological and paleoenvironmental constraits for the Messinian salinity crisis in the Mediterranean[END_REF][START_REF] Rouchy | Paleoenvironmental changes at the Messinian-Pliocene boundary in the eastern Mediterranean (southern Cyprus basins): significance of the Messinian Lago-Mare[END_REF] (see chapter 1 and 2 for further detail on Upper Gypsum deposits). 2022)) in both δ 18 OSO4 and δ 34 SSO4 recorded in phase 3 gypsum deposits than in PLG marginal basins (grey; [START_REF] Aloisi | The geochemical riddle of "low-salinity gypsum" deposits[END_REF][START_REF] Evans | Coupled measurements of delta O-18 and delta D of hydration water and salinity of fluid inclusions in gypsum from the Messinian Yesares Member, Sorbas Basin (SE Spain)[END_REF]; García-Veigas between the Miocene and Pliocene (Hardie and Lowenstein, 2004;Hsu and Montadert, 1978;[START_REF] Kastens | Proc. ODP, Init. Repts[END_REF]Ryan et al., 1973) (Figure 5. 4).

These cores penetrated the uppermost part of Messinian salinity crisis evaporites, corresponding to calcium sulfate deposits (gypsum or anhydrite) and marls alternations lying directly below Plio-Quaternary, mainly pelagic marine deposits and above halite formed during the preceding phase of the MSC (Figure 5. 4). The sites were chosen along an East-West Mediterranean transect: two sites to the west of the Sicily sill (site 134, West Sardinia and site 654, Tyrrhenian Sea) and two sites to the east of the Sicily sill, (Site 374, ionian abyssal plain and site 376, Florence Rise, west Cyprus) (Figure 5. 4).

In these cores, the intervals constituted of calcium sulfate minerals consist of various facies (i) laminated cumulate gypsum interpreted to have nucleated within a supersaturated water column (ii) clastic gypsum interpreted to originate from the resedimentation of gypsum formed upslope (iii) anhydrite interpreted to be the result of the dehydration of gypsum (iv) porphyroblastic gypsum interpreted to have formed by the rehydration of anhydrite [START_REF] Lugli | The deep record of the Messinian salinity crisis: evidence of a non-desiccated Mediterranean Sea[END_REF] (Figure 5. 4) (see section 2.2.1 for a detailed description of these facies).

Calcium sulfate deposits contain little amounts of iron sulfide (between 9 and 1440ppm, average of 235±352 ppm) (see Chapter 4). Although some of the calcium sulfate deposits investigated here have a clastic origin, we interpret the iron sulfide preserved in it as authigenic because of their small sizes and the high degree of reactivity of

Oxygen isotope of sulfate and multiple sulfur isotopes of sulfide and sulfate

The isotopic composition of sulfur and oxygen isotopes is expressed using the standard delta notation where R is the ratio between heavy and light isotope (RS= x S/ 32 S where x=34 or 33 and Ro= 18 O/ 16 O) and CDT and SMOW corresponds to the reference material used for calculation sulfur and oxygen isotopic composition respectively:

𝛿 "# 𝑆 % = ' )* + , )* +-./01 -14 × 1000 (1)
𝛿 "" 𝑆 % = '

)) + ,

)) + ,./01 -14 × 1000

(2)

𝛿 78 𝑂 % = ' :; + < :; + <.=>?@ -14 × 1000 (3)
When studying sulfur isotopes from the sedimentary record, a common approach is to present the difference in isotope composition between extreme oxidation state of sulfur that are most often conserved and available for measurements : sulfate (oxidation state +VI) (in the form of anhydrite, gypsum, barite or CAS) and sulfide (oxidation state -I or -II) (mostly in the form of pyrite or iron monosulfide) [START_REF] Findlay | Sulfide oxidation affects the preservation of sulfur isotope signals[END_REF]Johnston et al., 2005;Sansjofre et al., 2016b;Tostevin et al., 2014). The fractionation factors calculated from the isotopic composition of coeval sulfate and sulfide is the net effect of the numerous processes that have led to the formation and preservation of sulfur bearing minerals, all characterized by their own fractionation factors. The following formulas allow the calculation of net fractionation factors for both 33 S/ 32 S and 34 S/ 32 S isotopic ratios:

x αnet = RFeS/RSO4 (4)

Chapter 5-: Unraveling the roles of sulfide oxidation and of variable fractionation factors during microbial sulfate reduction ___________________________________________________________________________________________

221

"# Ɛ BCD = ( "# 𝛼 BCD -1) × 1000 ( 5)

"" 𝜆 BCD = IJ( )) K LMN ) IJ( )* K LMN ) (6)
The redox processes of the biogeochemical S cycle (e.g, MSR, sulfide reoxidation, disproportionation) often partition the 33 S isotopes (relatively to 32 S isotopes) differently from what is predicted by mass-dependent thermodynamic equilibrium (Ono et al., 2006). When this occurs, this phenomenon is recorded in the value of the fractionation factor λnet (equation 6, above) that will differ from the value predicted from thermodynamic equilibrium (λeq=0.515; Ono et al. (2006)). The study of net fractionation factors introduces a tool to investigate the potential occurrence and the various inflcuence of the above-mentioned redox processes.

The deviation of minor isotopes from thermodynamic equilibrium is also noted as follow for individual sulfur species:

∆ "" 𝑆 % = P'

)) + Q

)) + /01 -14 × 1000R -S'

)* + Q )* + /01 -14 T MU × 1000V (7)
In a δ 34 S-∆ 33 S diagram the relative data points position differ if the measured sulfur pool is the result of different sources of sulfur or if it is the result of a single specific redox process that fractionate isotopes.

Material and Methods

Gypsum and marls/mudstone samples interbeds were grinded and reacted with reductive solutions to transform them into Ag2S following the same method than Guibourdenche et al.

(2022) in order to measure 32 [START_REF] Waldeck | Calibrating the triple oxygen isotope composition of evaporite minerals as a proxy for marine sulfate[END_REF], as well as from that of PLG gypsum formed within Mediterranean marginal basins during the first phase of the MSC (δ 18 O=9.5±1‰, n=300) [START_REF] Aloisi | The geochemical riddle of "low-salinity gypsum" deposits[END_REF][START_REF] Evans | Coupled measurements of delta O-18 and delta D of hydration water and salinity of fluid inclusions in gypsum from the Messinian Yesares Member, Sorbas Basin (SE Spain)[END_REF]García-Veigas et al., 2018;Lu et al., 2001;[START_REF] Waldeck | Calibrating the triple oxygen isotope composition of evaporite minerals as a proxy for marine sulfate[END_REF] (Figure 5. 8).

Figure 5. 8 Comparison between δ 18 O values measured in this study at individual sites and values from the literature, from PLG various basins (pink) and from UG (green).

The values measured for the sites in the western Mediterranean (δ 18 O=13.3±0.3‰ at site 134 western Sardinia and δ 18 O=13.5±1.4‰ at site 654 in the Tyrrhenian basin) are included within the range previously measured for upper gypsum in western basins of the Mediterranean (δ 18 O=13.7±1.7‰, n=70) [START_REF] Aloisi | The geochemical riddle of "low-salinity gypsum" deposits[END_REF]García-Veigas et al., 2018;Lu et al., 2001;Waldeck MSC ocean (1) Sorbas basin (1,2,[START_REF] Lein | Isotopic composition of sulfur compounds in sediments of the Pacific[END_REF]4) Bajo Segura basin (4) Mallorca basin (4)

Vena del Gesso basin ( 5 extracted from calcium sulfate, there does not seem to be a significant dependence of the isotopic composition on the considered sites or facies in our dataset (Figure 5. 9).

Discussion

The main goal of the following discussion is to identify and quantify the relative importance of the various physical and biogeochemical processes in shaping the isotopic composition of iron sulfide preserved in deep basins of the Mediterranean. However, the calcium sulfate lithofacies that host these iron sulfides do not always corresponds to pristine primary gypsum samples but sometimes to anhydrite and porphyroblastic gypsum, for which the initial facies before dehydration/rehydration cycles is difficult to assess (Hardie and Lowenstein, 2004;[START_REF] Lugli | The deep record of the Messinian salinity crisis: evidence of a non-desiccated Mediterranean Sea[END_REF] ,or clastic gypsum possibly inherited from the dismantlement of gypsum deposits formed in a previous phase of the MSC (Hardie and Lowenstein, 2004;[START_REF] Lugli | The deep record of the Messinian salinity crisis: evidence of a non-desiccated Mediterranean Sea[END_REF]Roveri et al., 2014). Thus, it is important to understand the timing of precipitation of the calcium sulfate minerals we study. This question is crucial to tackle because the iron sulfide of during the deposition of phase 3 or (ii) are inherited from a previous stage of the MSC (iii) or reflect post depositional processes.

Origin of the sulfate ion constituting gypsum

The δ 34 S of UG calcium sulfate deposits is close to that of oceanic Messinian sulfate and to PLG gypsum after correction for gypsum precipitation [START_REF] Raab | Sulfur isotopic variations during seawater evaporation with fractional crystallization[END_REF]Thode and Monster, 1965;[START_REF] Van Driessche | Unraveling the Sulfate Sources of (Giant) Gypsum Crystals Using Gypsum Isotope Fractionation Factors[END_REF]. This suggests either marine waters or PLG gypsum or both as a potential sources of sulfate for UG formation (Figure 5. 6).

However, the δ 18 O of calcium sulfate minerals measured in this study is significantly higher than sulfate from both Messinian ocean and PLG gypsum (Figure 5. 8).

Sulfate oxygen does not exchange abiotically with water oxygen on time scales and pH relevant for this study, and must be mixed or cycled biogeochemically in its dissolved form to be modified (Lloyd, 1968;[START_REF] Turchyn | Oxygen isotopic composition of sulfate in deep sea pore fluid: evidence for rapid sulfur cycling[END_REF]. This means thats regardless of the facies of calcium sulfate minerals considered here (anhydrite, laminated, clastic or porphyroblastic gypsum), their isotope composition does not reflect the isotopic composition of older, resedimented, clastic evaporites (i.e. PLG gypsum), but that of dissolved sulfate that evolved isotopically through mixing or biogeochemical cycling. This does not preclude that PLG sulfate was a source of sulfate for UG formation but, if this is the case, it means that PLG sulfate must have been dissolved and biogeochemically processed before being incorporated into UG.

In addition: (i) laminated gypsum precipitated in the water column and is a primary gypsum facies [START_REF] Lugli | The deep record of the Messinian salinity crisis: evidence of a non-desiccated Mediterranean Sea[END_REF], (ii) clastic gypsum has likely experienced transport along small distances because it contains iron sulfides that are particularly sensitive to oxidation when does not cause significative changes in sulfate isotopic composition [START_REF] Pierre | The isotopic record of gypsum diagenesis in diluted solutions: Observations in natural salinas and experiments[END_REF]. We therefore interpret, regardless of the facies considered, that the isotopic compositions of the calcium sulfate minerals studied here reflect those of the parent water body in each basins at the end of phase 3 of the MSC and therefore that iron sulfide entrapped inside it were also formed during MSC phase 3. We will not develop further on the sources of dissolved sulfate for Upper Gypsum formation, since it is beyond the scope of this study and the isotopic composition of sulfate is relatively poorly sensitive to variations in the sources of the water masses.

Controls on the isotopic composition of iron sulfide

Together with the high δ 18 O of calcium sulfate minerals (Figure 5. 8), the fact that they contain iron sulfides-that form through the reaction of reactive iron with MSR products indicates that sulfur was biogechemically cycled in the 4 basins where these S-bearing minerals formed.

There are different processes that have the potential to shape the isotopic composition of iron sulfide minerals. These include : sulfide pooling during the consumption of sulfate in a restricted reservoir (Houghton et al., 2022;[START_REF] Masterson | Interpreting multiple sulfur isotope signals in modern anoxic sediments using a full diagenetic model (California-Mexico margin: Alfonso Basin)[END_REF]Pasquier et al., 2021b;Pasquier et al., 2017), occurrence of re-oxidative processes [START_REF] Findlay | Sulfide oxidation affects the preservation of sulfur isotope signals[END_REF]Pellerin et al., 2015;Pellerin et al., 2019), variation in the fractionation factors during microbial sulfate reduction (Jiang et al., 2022;Leavitt et al., 2013;Mengdi et al., 2022) or mixing of sulfide pools having different isotopic composition due to their formation in different biogeochemical conditions [START_REF] Fike | Rethinking the ancient sulfur cycle[END_REF]Pasquier et al., 2021b). These biogeochemical processes can occur concomitantly or not and their prevalence are controlled by environmental settings during iron sulfide formation. In the following sections, we modelled 4 biogeochemical In all the model scenarios, we set the δ 34 S-∆ 33 S of the initial reacting sulfate processed by MSR to that of average Messinian sea water (Masterson et al., 2016), since we demonstrated in the previous section that the sulfate isotopic composition of the Mediterranean basins considered had a similar isotopic composition.

5.8.1 First scenario: Pore water sulfur isotopic distillation as a driver for anomalously high isotopic composition of iron sulfides?

Large amplitude oscillations in the isotopic composition of sedimentary sulfide between 34 Senriched solid sulfides and 34 S-depleted solid sulfide similar to what is observed in our samples are often linked to variation in the pooling of sulfide, driven by changes in sedimentary settings due to water level oscillations (Houghton et al., 2022;Pasquier et al., 2017). In these studies, the production of 34 S-enriched sulfide deviates from what is expected for hydrogen sulfide produced by MSR in classic sedimentary environments (i.e. between εMSR= -70 and -30‰ (e.g. Masterson et al., 2016;[START_REF] Prince | Evaporite deposition in the mid-Neoproterozoic as a driver for changes in seawater chemistry and the biogeochemical cycle of sulfur[END_REF]Tostevin et al., 2014)). For example, based on a correlation between δ 34 Spyrite and sedimentation rates, some authors have interpreted that the isotopic composition of pyrite is sometimes driven by the relative restriction of pore water, itself controlled by fluctuation in the sedimentation rate. During periods of high sedimentation rates, the efficient disconnection of porewater sulfate from the overlying sulfate rich water column allowed the efficient distillation of the sulfate and the preservation of δ 34 Spyrite enriched isotopic compositions (Pasquier et al., 2017). In the modern Peru margin, it is the fluctuation in organic matter delivery linked to cyclic contraction and This phenomenon occurs when the rate at which sulfate is removed from the system by MSR overcomes the rate at which it is replenished by diffusion from the water column and can be the result of a range of environmental controls.

In order to test the potential influence of sulfide pooling on iron sulfide isotopic composition, we set our model to simulate the progressive enrichment of sulfide due to an isotopic Rayleigh distillation during the consumption of sulfate and the production of sulfide by microbial sulfate reduction. In addition, the porewater reservoir in which sulfate distillation occurs, is rarely completely disconnected from the overlying water column that represent a virtually infinite reservoir of sulfate. We thus mixed the sulfide resulting from Rayleigh distillation to a pool of undistilled, 34 S-depleted sulfide. This model simulation could correspond to the mixing We applied a Rayleigh type distillation to sulfur isotopes of the reactant and product of microbial sulfate reduction in order to simulate the effect of MSR in a close system where hydrogen sulfide accumulates at the expense of sulfate consumption. Considering that the isotopic fractionation during the formation of iron sulfide is negligible compared to the large isotopic excursion produced by microbial induced processes (Eldridge et al., 2021), we approximated that the isotopic composition of iron sulfide was equivalent to the isotopic composition of hydrogen sulfide produced by MSR (Figure 5. 10). We used the following Rayleigh distillation equations to calculate the isotopic ratios of both the reacting sulfate 3x/32 R SO4 = 3x S/ 32 S(SO4) and the cumulated sulfide 3x/32 RFeS = 3x S/ 32 S(FeS) with

x corresponding either to 33 or 34 isotopes.

3x/32 R SO4 (f SO4 ) = 3x/32 R SO4 (f SO4 =1)*f SO4 (αMSR-1) (A) 3x/32 R FeS (f SO4 ) = ( 3x/32 R SO4 (f SO4 =1) -3x/32 R SO4 (f SO4 )*f SO4 ) /(1-f SO4 ) (B)
In equations (A) and (B), fSO4 represents the pool of sulfate that remains, it varies between 0 and 1, when fSO4=0 the sulfate pool is completely exhausted, all the sulfur in the system corresponds to sulfide whereas fSO4=1 corresponds to an undistilled sulfide pool (Figure 5. 10).

We modelled this sulfide pooling for different fractionation factors that are realistic to occur during microbial sulfate reduction, with αMSR=0.93, 0.95 and 0.98 representing εMSR=-70‰, -50‰ and -20‰ respectively.

We recalculated the isotopic composition for the iron sulfide modelled isotopic ratios ( 3x/32 RFeS with x=33 or 34) using equation ( 1) and ( 2) to obtain δ 34 Spooled and δ 33 Spooled.

We then modelled the mixing between the sulfide affected by sulfide pooling (δ 34 Spooled and δ 33 Spooled) and an undistilled reservoir of sulfide (δ 34 Sundist and δ 33 Sundist) produced by MSR with the various αMSR mentioned above, and using a starting sulfate having an oceanic isotopic composition.

The result of this mixing was calculated using the following equation: This agrees with the relatively large TOC reported at some sites for non-evaporitic interbeds (up to 5.3%) reported in IODP/DSDP reports (Deroo, 1978;McIver and RD, 1973;Pierre and Rouchy, 1990b;Sigl et al.). High influxes of organic carbon to the MSR zone could have modulated the pore-water distillation of sulfur isotopes through high sulfate consumption by MSR. The occurrence of sulfide pooling seems a realistic hypothesis to explain the isotopic composition of iron sulfides formed in non-evaporitic interbeds in which we can't estimate the sulfate concentration but that bore pore waters that were most probably undersaturated with respect to gypsum.

𝛿 "W 𝑆 X7 = 𝛿 "W 𝑆 YZZ[C\ × 𝑓 ^%W + 𝛿 "W 𝑆 `B\%-D × (1 -𝑓 ^%W ) (C)
Conversely, if the scenario exposed in Panel A -Figure 5. 11 holds true, it implies that iron sulfides were formed in porewaters where the original sulfate reservoir was nearly completely (95 to 99%) consumed by MSR. Indeed, more than 50% of the sulfides measured in gypsum plot in the area defined by fSO4= 0.05 to 0.01.

This result is at odds with the fact that iron sulfides occur within gypsum deposits. Indeed, to produce such an enrichment by isotopic distillation, the sulfate reservoir where MSR occurred needs to be completely disconnected from any source of sulfate that would continuously replenish it (Figure 5. 11). However, given that iron sulfide are preserved within calcium sulfate evaporites that (i) have the isotopic composition of contemporaneous sea water and

(ii) are higly soluble, the porewater, was at minima, at gypsum saturation during iron sulfide formation (Hoareau et al., 2011).

Chapter 5-: Unraveling the roles of sulfide oxidation and of variable fractionation factors during microbial sulfate reduction ___________________________________________________________________________________________ 237

The surrounding gypsum therefore possibly constituted a virtually infinite reservoir of sulfate within the zone of formation of iron sulfide. Therefore, the 95 to 99% of sulfate reservoir consumption required to explain most of our data in a scenario where Rayleigh distillation control sulfide isotopic composition is highly unlikely. This is further supported by the inverse correlation between δ 34 S isotopic composition and iron sulfide concentration, with the lowest concentrations recorded in samples where the heavier isotopic composition were measured (Figure 5. 12). Indeed, one could expect that when sulfide isotope variations are controlled by a Rayleigh distillation, the enrichment of sulfide isotopes in 34 S isotopes is a function of the flux of sulfide produced by MSR and preserved in a solid form, therefore, the inverse correlation than what is observed in our dataset should be expected if this scenario was true.

Scenario 2: concurrent influence of variable fractionation factors and sulfide pooling

Distillation of sulfur isotopes is not the only process that can produce 34 S rich iron sulfide, variations in the fractionation factor expressed during MSR is sometimes evoked. As mentioned above, fluctuations in the rate of microbial sulfate consumption have been inferred to control the isotopic composition of pyrite preserved in sediment underlying the Peruvian oxygen minimum zone (OMZ) by controlling the extent of sulfide pooling in sedimentary pore waters (Pasquier et al., 2021b). In addition, another important control on the rate of sulfate consumption by MSR is the cell specific rate of sulfate reduction, csSRR that is also a major control on the extent of the isotopic fractionation expressed during MSR (Leavitt et al., 2013;Sim et al., 2011a;Sim et al., 2011b;[START_REF] Wing | Intracellular metabolite levels shape sulfur isotope fractionation during microbial sulfate respiration[END_REF] sulfide produced) for high csSRR (Leavitt et al., 2013;Sim et al., 2011a;Sim et al., 2011b).

Therefore, sulfide pooling and varying fractionation factors during MSR possibly driven by fluctuating csSRR could be two concomitant controls on the isotopic composition of iron sulfide. We tested the hypothesis of their combined influences in a restricted reservoir of sulfate,

where MSR allow for a significant consumption of sulfate with varying fractionation factors, due to fluctuations in environmental parameters (Figure 5. 13).

To express the variability of fractionation factors possibly taking place during microbial sulfate reduction we used the nonlinear regression model from Leavitt et al. (2013) We rearranged equations (D) and (E) to obtain the relationship that link 34 Ɛ and 33 λ to normalized csSRR.

𝑐𝑠𝑆𝑅𝑅 = IJ( )) m >=e f )* m >QL )* m >gh f )* m >QL ) i )) j k,l (
"" 𝜆 ^-n = (𝑒 ip-X++× )) j k,l (𝜆 ^qW -𝜆 ^%B ) + 𝜆 ^%B (F)

rr 𝛼 ^-n = ( rs 𝛼 ^-n ) rr T k,l (G)
In our model we thus explore the entire range of possible isotopic fractionation produced by MSR by varying 34 ƐMSR between -5 and -70‰ (Figure 5. 13).

We then modelled the production of sulfide in a restricted reservoir by using equation (A) and as thin crust on other minerals (see chapter 4). Given that most of these grains corresponded to gypsum that is highly soluble, they represented a virtually infinite pool of sulfate available for MSR (Hoareau et al., 2011). This precludes sulfide pooling to be a major control on the isotopic composition of iron sulfide that were extracted from calcium sulfate layers in this study. We conclude from both scenario 1 and 2 that unpooled oceanic sulfate was most probably the original reservoir of sulfur used by microbial sulfate reducers to produce hydrogen sulfide.

Scenario 3: influence of variable sulfur fractionation factors in a virtually infinite pool of sulfate

The results of the two previous model exercises confirmed the formation of iron sulfide within evaporitic layers in a virtually infinite reservoir of sulfate, being either the bottom water or the sediment porewater. In the MSR zone of this sulfate saturated environment, the rate of microbial sulfate reduction could have been heterogeneous, resulting in variations of the fractionation factors expressed during MSR with depth within MSR zone. For example, the top of the MSR zone where fresh organic matter is delivered by diffusion from the overlying water column or porewater environment could have yielded higher rates of microbial sulfate reduction than deeper horizon with access to fewer amount and/or more refractory organic matter. Lower rates at the bottom of the MSR zone would lead to large fractionation factors (close to 34 ƐMSR ≈-70‰) whereas higher rates at the surface tend to diminish the extent of isotopic fractionation. Diffusion, bioturbation or reworking events could lead these two isotopically distinct reservoirs to mix occasionally or on a regular basis, leading to the modification of the isotopic composition of preserved sulfide from the initial one produced in the separate horizons of the MSR zone. We tested this scenario by considering the mixing between different pools of sulfide produced by MSR, the first one exploring the entire range possibly expressed during MSR and the second one produced by MSR with a fractionation at equilibrium between sulfate and sulfide (Figure 5. 15).

We used equation (F) and (G) to simulate the dependence of 33 S/ 32 S and 34 S/ 32 S partitioning to csSRR. These equations were rearranged with equation ( 4) to obtain the sulfide isotopic ratios of iron sulfide as a function of the extent of MSR fractionation 3x/32 R FeS ( 34 ƐMSR) (with x=33 or 34). The obtained isotopic ratios where then used to calculate the isotopic composition of iron sulfide produced with various fractionation factors (δ 34 SƐMSR and δ 33 SƐMSR)

Chapter 5-: Unraveling the roles of sulfide oxidation and of variable fractionation factors during microbial sulfate reduction ___________________________________________________________________________________________ 245 using equations ( 1) and ( 2). This pool of sulfide was then mixed with a pool of sulfide produced with the maximum fractionation factor of -70‰, (δ 34 SƐMSR=-70 and δ 33 SƐMSR=-70) using equation

(C).
The results of this model scenario are shown in Figure 5. 16 together with the values measured in our samples. The iron sulfide preserved in our samples can be explained by mixing various proportion of sulfide produced with different fractionation factors. highly favorable conditions for microbial sulfate reducers [START_REF] Bradley | Patterns of sulfur isotope fractionation during microbial sulfate reduction[END_REF]. Further, given the higher TOC content reported in non-evaporitic interbeds than observed in calcium sulfate layers, it is likely that if organic matter delivery controlled the fractionation during MSR in our samples, iron sulfide preserved in marls and mudstone would display higher δ 34 S than the one preserved in evaporitic layers. However, the iron sulfides that have anomalously high δ 34 S are mostly preserved in gypsum, whereas marls generally display isotopic composition reflective of higher fractionation (Figure 5. 16). However, even though they are rarely observed in natural environments, some of the environments where such small fractionation is observed correspond to anomalo-haline water bodies with relatively large concentrations of sulfate, (Avrahamov et al., 2014;Mengdi et al., 2022), probably similar to the conditions under which our samples formed.

Given that sulfide pooling is not a valid hypothesis to explain the enrichment in heavy isotopes by several other authors in the literature [START_REF] Aloisi | The geochemical riddle of "low-salinity gypsum" deposits[END_REF][START_REF] Waldeck | Calibrating the triple oxygen isotope composition of evaporite minerals as a proxy for marine sulfate[END_REF]. We detail why in the following paragraph.

MSR preferentially metabolizes the light isotopes of sulfur and oxygen ( 32 S and 16 O respectively) and results in the enrichment of the remaining sulfate in heavy isotopes ( 34 S and 18 O respectively) (Brunner and Bernasconi, 2005). This leads to a co-evolution of δ 34 S and δ 18 O along lines having a gradient of slope between 0.5 and 3.5 [START_REF] Fotherby | Modelling the Effects of Non-Steady State Transport Dynamics on the Sulfur and Oxygen Isotope Composition of Sulfate in Sedimentary Pore Fluids[END_REF] (see grey lines in Figure 5. 17). However, in our dataset, the δ 18 O offset from sea water sulfate is not accompanied by the expected concomitant substantial increase in δ 34 S (Brunner and Bernasconi, 2005) (Figure 5. 17). This type of δ 34 S -δ 18 O distribution is generally associated to cryptic sulfur cycling, i.e., the occurrence of a quantitative reoxidation of sulfide into sulfate [START_REF] Antler | The Sedimentary Carbon-Sulfur-Iron Interplay -A Lesson From East Anglian Salt Marsh Sediments[END_REF][START_REF] Johnston | Placing an upper limit on cryptic marine sulphur cycling[END_REF][START_REF] Mills | Geochemical evidence for cryptic sulfur cycling in salt marsh sediments[END_REF]. This phenomenon leads to the reinjection of the MSR sulfur product ( 34 S-depleted) into the residual reactant sulfate pool ( 34 S-enriched). This results in no net effect of sulfur cycling on both sulfur species (sulfide and sulfate) concentrations and δ 34 S isotopic composition. Conversely, the oxygen from sulfate tends to equilibrate isotopically with ambient water when cycled biogeochemically (either by MSR or re-oxidative processes) [START_REF] Antler | The Sedimentary Carbon-Sulfur-Iron Interplay -A Lesson From East Anglian Salt Marsh Sediments[END_REF][START_REF] Antler | Coupled sulfur and oxygen isotope insight into bacterial sulfate reduction in the natural environment[END_REF]Brunner et al., 2005;[START_REF] Mills | Geochemical evidence for cryptic sulfur cycling in salt marsh sediments[END_REF]. Its composition will stop to increase when offset from ambient water oxygen of ∼23‰ that represent the isotopic equilibrium between intracellular sulfite and water (Bertran et al., 2020;Brunner et al., 2005;Zeebe, 2010). The occurrence of sulfide reoxidation seems to be a widespread phenomenon during the phase 3 of the MSC, since both marginal and deep, eastern and western basins of the Mediterranean present δ 34 S values that are relatively constant and centered on the messinian ocean isotopic composition (Figure 5. [START_REF] Aloisi | The geochemical riddle of "low-salinity gypsum" deposits[END_REF]García-Veigas et al., 2018;Pierre and Fontes, 1982) and those from deep basins of the Mediterranean measured in this study. The possible evolution of the isotopic composition of residual sulfate during MSR is indicated by the grey lines [START_REF] Fotherby | Modelling the Effects of Non-Steady State Transport Dynamics on the Sulfur and Oxygen Isotope Composition of Sulfate in Sedimentary Pore Fluids[END_REF].

However, whether or not this re-oxidative cycling of sulfur recorded in gypsum, and therefore possibly occurring far from the MSR zone where iron sulfide have been produced remains unclear. To tackle this we calculated net fractionation factors εnet and 33 λnet following equation ( 5) and ( 6) between sulfide produced and Messinian ocean. We didn't directly used coeval sulfate measurement from the calcium sulfate minerals entrapping iron sulfide because both minerals have potentially been formed at different depths within the basin. However, given that we demonstrated that iron sulfate formed in an environment saturated with sulfate, that We then modelled the evolution of the isotopic composition of the hydrogen sulfide reservoir that vary in size not because it is not produced by MSR but because it got consumed by reoxidative process using the following equation.

3x/32 R FeS (f ox )= 3x/32 R FeS ( 34 ƐMSR) x f ox (αox-1) 3x/32 R FeS (f ox ) corresponds to the isotopic composition of the remaining sulfide, initially produced by MSR with variable fractionation factors and then progressively consumed by sulfide reoxidation. αox is the fractionation factor associated with reoxidation of sulfide and f ox the amount of sulfide remaining in the system, with f ox =1 meaning that no H2S has been consumed by the considered re-oxidative process and f ox =0 meaning that all of the H2S produced by MSR have been consumed by a re-oxidative process. Every scenario of sulfide reoxidation allows to explain the wide range of variations expressed in our samples. In general, the models results predict higher δ 34 S of iron sulfide when the amount of sulfide being reoxidized is larger and the fractionation factor during MSR is smaller (Figure 5. 20). This could provide a satisfying explanation for the anticorrelation between iron sulfide concentration and their isotopic signature (Figure 5. 12), with lesser amount of sulfide being preserved when δ 34 S are higher because more reoxidation of sulfide occurred before it got scavenged by reactive iron to form iron sulfide.

Although the estimated fractionation factors and fraction of sulfide being reoxidized necessary to produce the most 34 S-enriched iron sulfide measured in this study vary between re-oxidative scenarios, in all scenarios, these values are explained only if the great majority of sulfide produced by MSR was reoxidized. With an amount of sulfide being reoxidized of 90 to 99% for abiotic reoxidation, 70 to 99.5% for phototrophic oxidation and 99.9 to 99.998% for chemolithotrophic reoxidation necessary order to explain the isotopic composition of the most enriched sulfide of measured in this study (Figure 5. 20). This agrees with the hypothesis of a cryptic cycling of sulfur during the deposition of calcium sulfate minerals, as previously suggested by the oxygen isotopic composition of sulfate.

The exact nature of the re-oxidative process at play is difficult to identify since only a few studies, have assessed the fractionation factors associated with sulfur oxidation (Johnston et al., 2005;[START_REF] Zerkle | Fractionation of multiple sulfur isotopes during phototrophic oxidation of sulfide and elemental sulfur by a green sulfur bacterium[END_REF]Zerkle et al., 2016). Further, these studies did not explore all the deposition of Upper Gypsum, we confirmed that the enrichment in the isotopic composition of iron sulfide is not always only driven by the distillation of isotopes and accumulation of sulfide in restricted reservoirs (i.e sulfide pooling) as suggested in a large range of recent studies (e.g., [START_REF] Jørgensen | Sulfur biogeochemical cycle of marine sediments[END_REF]Jørgensen et al., 2019;[START_REF] Masterson | Interpreting multiple sulfur isotope signals in modern anoxic sediments using a full diagenetic model (California-Mexico margin: Alfonso Basin)[END_REF]Pasquier et al., 2021b). As shown in this study for the very specific case of MSC phase 3 deposits, it can be also controlled to a certain extent by the variability in the fractionation factors 34 ƐMSR and 33 λMSR expressed during MSR. In models that use sulfur isotopic compositions to reconstruct past sulfur cycle, a relatively arbitrarily fixed 34 εMSR value (that can vary between -30 and -70‰), is often assumed (e.g. Kunzmann et al., 2017;Masterson et al., 2016;[START_REF] Prince | Evaporite deposition in the mid-Neoproterozoic as a driver for changes in seawater chemistry and the biogeochemical cycle of sulfur[END_REF]Saitoh et al., 2021;Thomazo et al., 2019). Discriminating the effect of sulfide pooling from changes in the magnitude of isotope fractionation is of primary importance because in past reconstruction of the sulfur cycle, isotope-based estimates of sulfur fluxes depend on the fractionation factor(s) used to model the biogeochemical processes into play (Kunzmann et al., 2017;Sansjofre et al., 2016a;Tostevin et al., 2014). As shown in Figure 5. 2 the calculated amount of sulfate depletion greatly depends on the fractionation factor choosen. A better assessment on the actual range of fractionation possibly expressed in the environment as well as a better knowledge on the environmental settings that controls it, is thus needed to refine our understanding of sulfur cycle in past and present environments. In conclusion, the role of varying fractionation factors in shaping the isotopic composition of sulfide produced by MSR should be considered when trying to extract paleoenvironmental informations from these complex archives.

Conclusion

In this work, we also demonstrated the occurrence of a very active re-oxidative cycling of sulfur that allowed the great majority of the sulfide produced by MSR to be reoxidized into sulfate. Whereas we have shown that this phenomenon was occurring in a marginal restricted the MSC. In opposition to PLG gypsum, the gypsum sulfate we measured here presented a significant deviation from the Messinian Sea water sulfate isotopic composition. This could be due either to (i) a larger degree of reequilibration of the oxygen from the sulfate with the ambient water oxygen isotopic composition than during the PLG that have values centered around the MSC ocean sulfate isotopic composition or (ii) a similar degree isotopic equilibration with water than during PLG formation but with a water that had a heavier δ 18 OH2O isotopic composition, possibly due to the fact that the upper gypsum parent water might have underwent a larger degree of evaporation. In addition, the significant difference in the δ 18 OSO4 sulfate isotopic composition between the sites located in the western and eastern basins of the Mediterranean suggest that these two areas evolved biogeochemically independently from each other during the formation of MSC phase 3 gypsum. The Sicily sill, separating the western Mediterranean from the eastern Mediterranean, lies nowadays at 430m depth and is understood to not have been greatly uplifted since the Messinian (Blanc, 2006). A base level fall during the MSC might have disconnected the western from the eastern Mediterranean basins [START_REF] Andreetto | Freshening of the Mediterranean Salt Giant: controversies and certainties around the terminal (Upper Gypsum and Lago-Mare) phases of the Messinian salinity crisis[END_REF][START_REF] Lofi | Erosional processes and paleo-environmental changes in the Western Gulf of Lions (SW France) during the Messinian Salinity Crisis[END_REF][START_REF] Micallef | Evidence of the Zanclean megaflood in the eastern Mediterranean Basin[END_REF]. This further support the results presented in the Chapter 4 of this manuscript. More data from the upper gypsum in the eastern basins are however needed to corroborate the separated biogeochemical evolution of eastern and western basins of the Mediterranean toward the end of the MSC.

Finally, the occurrence of the complete reoxidation of sulfide produced by MSR, i.e., cryptic sulfur cycling, and in general any re-oxidative cycling of sulfur is most of the time overlooked, especially in sedimentary archives. However, this has great implication on our understanding

Chapter 5-: Unraveling the roles of sulfide oxidation and of variable fractionation factors during microbial sulfate reduction ___________________________________________________________________________________________ 257 of the cycling of sulfur and its links with the cycling of major element within the oceanic realm.

Indeed, in most studies that investigate organic carbon recycling within marine sediments, porewater sulfate concentrations or radiolabeling of sulfate are two methods often used to infer the fluxes of carbon being oxidized through microbial sulfate reduction [START_REF] Bowles | Global rates of marine sulfate reduction and implications for sub-sea-floor metabolic activities[END_REF][START_REF] Bradley | Widespread energy limitation to life in global subseafloor sediments[END_REF]Egger et al., 2018;[START_REF] Jørgensen | Sulfur biogeochemical cycle of marine sediments[END_REF]. Whereas this approach allows to quantify the net fluxes of hydrogen sulfide produced by microbial sulfate reduction, rapid re-oxidation of sulfide back to sulfate (Canfield et al., 2010;[START_REF] Liu | Early diagenesis of iron and sulfur in Bornholm Basin sediments: The role of near-surface pyrite formation[END_REF][START_REF] Johnston | Placing an upper limit on cryptic marine sulphur cycling[END_REF][START_REF] Treude | Rapid sulfur cycling in sediments from the Peruvian oxygen minimum zone featuring simultaneous sulfate reduction and sulfide oxidation[END_REF] might minimize the gross amount of sulfide being actually produced by MSR in certain environments, and therefore lead to the underestimation of the actual fluxes of organic carbon being remineralized through MSR. Therefore, measurement of multiple sulfur isotope, together with a better assessment of the gross rates of MSR seems crucial to be further applied in surficial sediment and past archives in order to assess the impact of reoxidative sulfur cycling on major element cycling. In the absence of reaction rate measurements, multiple sulfur isotopes together with a numerical modelling approach appear to be one of the only tool available to detect and quantify the role of re-oxidative sulfur cycling in the sedimentary record.

Chapter 6-Conclusion and perspectives ___________________________________________________________________________________________ 259 Chapter 6. These lithological alternations are interpreted to reflect cyclical hydrological changes in Mediterranean basins (both marginal and deep), themselves controlled by orbitally driven climate oscillations (Andreetto et al., 2022;Lugli et al., 2010;[START_REF] Manzi | A new facies model for the Upper Gypsum of Sicily (Italy): chronological and paleoenvironmental constraits for the Messinian salinity crisis in the Mediterranean[END_REF]. Here we showed that these changes in the hydrology (from more humid during precession minima to more arid during precession maxima) also controlled the isotopic signature of the iron sulfide preserved in evaporitic deposits and marl interbeds, with iron sulfides having low δ 34 S during marl deposition and higher δ 34 S during gypsum deposition (Figure 6. 2). We interpreted these results using a numerical modelling approach and demonstrated that contrasting δ 34 S values correspond to drastically different sulfur redox cycling modes during periods of gypsum precipitation versus periods of marl deposition (Figure 6. 3Figure 6. 4). The high δ 34 S and low ∆ 33 S values that characterize iron sulfides within gypsum were (i) the result of microbial sulfate reduction expressing limited isotope fractionations (ii) produced in a period where the quasi totality of sulfide produced by MSR was subsequently re-oxidized, resulting in a cryptic cycling of sulfur (Figure 6. 3Figure 6. 4). In comparison, iron sulfides in marls yield on average lower δ 34 S and higher ∆ 33 S. This suggests that a less intense reoxidation and larger fractionation during microbial sulfate reduction characterized the biogeochemical sulfur cycle during the formation of marls (Figure 6. 3Figure 6. 4).

Conclusion and perspectives

In gypsum and marl interbeds formed at the onset of the MSC (PLG deposits), we interpreted these different trends to reflect the alternation between a cryptic cycling of sulfur during gypsum formation -involving the remineralization of comparable fluxes of benthic organic carbon through MSR than in the most productive oceanic regions of the world ocean -and overall smaller fluxes of organic carbon being remineralized by MSR and a less intense reoxidative sulfur cycle during marl formation under a stratified and intermittently euxinic water column (Figure 6. 5) (Damsté et al., 1995). The cryptic cycling of sulfur during gypsum precipitation was driven by a consortium of sulfur-using bacteria, including giant sulfide oxidizers, of which the remains are found in PLG gypsum (Figure 6. 5, panel C); these are functionally similar to the ones found on the sea floor of highly productive modern oceanic margins. The fact that more than 80% of the sulfate ultimately used in gypsum formation was first involved in this cryptic sulfur cycling supports their role as a trigger for gypsum important factor that control these environmental parameters (e.g., [START_REF] Fike | Rethinking the ancient sulfur cycle[END_REF]Houghton et al., 2022;Pasquier et al., 2021a;Pasquier et al., 2021b;Pasquier et al., 2017). Sea level exerts a major control on the isotopic composition of iron sulfides by affecting various environmental parameters such as sedimentation rates, benthic fluxes of organic carbon, fluxes of sulfide re-oxidation and sediment reworking (Aller, 2010;Aller, 2014;Dunne et al., 2007;[START_REF] Jørgensen | Sulfur biogeochemical cycle of marine sediments[END_REF]Middelburg et al., 1997). I built on these observations to provide insights on the water level of Mediterranean basins during the third and terminal stage of the MSC, that is a particularly controversial issue in the MSC literature. I coupled my sulfur isotopic approach to petrographic observations. This allowed me to find out that the iron sulfide entrapped in both gypsum and marl alternation during stage 3 mainly consist in early diagenetic nanometric iron sulfides globules (e.g., Figure 6. 6). In Figure 6. 7, it is possible to see that there is a relationship between the depth of deposition of early Given that the isotopic composition of these early diagenetic iron sulfide precipitates was particularly enriched ( -13±16‰) and variable (from -40 to +13‰), I suggested that the units that host these iron sulfides couldn't have deposited at the depth at which they are found nowadays (between 1965 and 4450 mbsl), and were formed under a water column that was -at maximum -1500m thick (Figure 6. 7).

Chapter 6-Conclusion and perspectives ___________________________________________________________________________________________ 267

In a collaboration with a PhD student colleague from the SaltGiant ETN, I used this maximal depth of deposition in a paleo-bathymetric reconstruction to constrain the amplitude of the base level fall that characterized the end of the Messinian Salinity Crisis. The result of this approach suggest a water level drop of 1100m in the western Mediterranean and a 2500m drawdown in the eastern Mediterranean, implying that these two basins were disconnected at the end of the MSC, and supporting the findings of studies using independent seismic markers (e.g., [START_REF] Camerlenghi | Uncovering the Mediterranean Salt Giant (MEDSALT)-Scientific Networking as Incubator of Cross-disciplinary Research in Earth Sciences[END_REF][START_REF] Andreetto | Freshening of the Mediterranean Salt Giant: controversies and certainties around the terminal (Upper Gypsum and Lago-Mare) phases of the Messinian salinity crisis[END_REF][START_REF] Lofi | Erosional processes and paleo-environmental changes in the Western Gulf of Lions (SW France) during the Messinian Salinity Crisis[END_REF][START_REF] Micallef | Evidence of the Zanclean megaflood in the eastern Mediterranean Basin[END_REF].

Overall, the results presented in this thesis demonstrate that the multiple isotopic composition of sulfur-bearing minerals has the potential to unravel the dynamics of sulfur cycling in the micro-environment where they formed. Specifically, this study allowed to highlight the importance of the variability of two controls on the isotopic composition of sedimentary sulfur: sulfide re-oxidation and fractionation factors during microbial sulfate reduction.

Importance of sulfide re-oxidation

The use of multiple sulfur isotopes has allowed to quantify the relative importance of the fluxes of sulfide being re-oxidised, compared to the fluxes of sulfate initially reduced to sulfide by MSR. In chapters 3 and 5 of this thesis, we showed that the overall importance of the reoxidative flux of sulfur can vary temporally depending on environmental conditions, whereas studies on past and present sediments often consider fixed fluxes of sulfide being re-oxidized (e.g. 90%) in marine sediments, based on estimates determined in few studies in shallow water sediments of the Baltic Sea (Fossing and Jørgensen, 1990;[START_REF] Jørgensen | A Thiosulfate Shunt in the Sulfur Cycle of Marine Sediments[END_REF][START_REF] Thamdrup | Thiosulfate and Sulfite Distributions in Porewater of Marine-Sediments Related to Manganese, Iron, and Sulfur Geochemistry[END_REF]. Further, intense sulfide re-oxidation in the uppermost part of surficial sediments causes a huge discrepancy between gross and net rates of microbial sulfate reduction, especially (e.g., [START_REF] Liu | Early diagenesis of iron and sulfur in Bornholm Basin sediments: The role of near-surface pyrite formation[END_REF]Moeslund et al., 1994). The different methods used to estimate the amount of organic carbon being remineralized by MSR rely, amongst other things, on the correct assessment of gross rates of MSR. The estimate of yearly MSR rates however yields huge discrepancies with almost one order of magnitude between the lowest and highest of these estimates (estimates of between 8.4 and 80 TmolS.yr -1 ) [START_REF] Bowles | Global rates of marine sulfate reduction and implications for sub-sea-floor metabolic activities[END_REF][START_REF] Bradley | Widespread energy limitation to life in global subseafloor sediments[END_REF]Canfield et al., 2005;Jørgensen and Kasten, 2006;Thullner et al., 2009). Therefore, a precise and more systematic assessment of the relative importance of reoxidative fluxes is needed to better constrain the overall importance of organic carbon remineralization by microbial sulfate reduction in the ocean.

Assessing the environmental controls on the relative amount of sulfide being reoxidised is also crucial because, in the context of global climate change, increased sea surface temperatures will reduce oceanic oxygen uptake and cause the upward displacement of the zone where hydrogen sulfide production starts (Shaffer et al., 2009). This phenomenon is accentuated in coastal environments, where dysoxic waters expand in response to the massive release of nutrients from agricultural activity. This leads to the eutrophication of surface waters, and to massive oxygen consumption by organic matter remineralization which overcome oxygen supply via ventilation and mixing (Breitburg et al., 2018;Diaz and Rosenberg, 2008;Kemp et al., 2009).

While several studies have shown that sulfide oxidizing bacteria have the potential to prevent the release of toxic hydrogen sulfide into the water column (Callbeck et al., 2021;[START_REF] Seitaj | Cable bacteria generate a firewall against euxinia in seasonally hypoxic basins[END_REF], it is critical to understand how the biogeochemical cycling of sulfur responds to these massive changes in oceanic biogeochemistry. In the absence of incubation experiments and/or analyses of pore water geochemical profiles, multiple sulfur isotope measurement can provide a valuable tool to assess the fluxes of sulfur involved in microbial sulfate and subsequent sulfide reoxidation in both modern and ancient sediments, as shown in this work.

Variable sulfur fractionation factor expressed during MSR

The range of sulfur isotope fractionation expressed in natural environments during MSR depends on a variety of factors such as the cell specific rate of microbial sulfate reduction (Habicht and Canfield, 2001;Leavitt et al., 2013;Sim et al., 2011a;Sim et al., 2011b), temperature, pH, type of strains, structure and size of the sulfate reducing community [START_REF] Brüchert | Controls on stable sulfur isotope fractionation during bacterial sulfate reduction in Arctic sediments[END_REF][START_REF] Colangelo-Lillis | Diversity decoupled from sulfur isotope fractionation in a sulfatereducing microbial community[END_REF][START_REF] Stam | 34S/32S fractionation by sulfate-reducing microbial communities in estuarine sediments[END_REF], nature and accessibility of organic substrates, environmental concentrations of sulfide and sulfate [START_REF] Bradley | Patterns of sulfur isotope fractionation during microbial sulfate reduction[END_REF][START_REF] Wing | Intracellular metabolite levels shape sulfur isotope fractionation during microbial sulfate respiration[END_REF], or even the mineralogy of the sediments (Lin et al., 2022).

Further, the range of fractionation that is potentially expressed in natural environments is, itself, debated (Jørgensen et al., 2019) (see section 3.4.3.2 for a detailed discussion on this issue). By using coeval sulfate-sulfide multiple sulfur isotopes measurement we highlighted As discussed in chapter 1, gateway changes provoked deep-water stagnation leading to an increase in bottom water residence time and oxygen depletion that greatly impacted the carbon biogeochemical cycle and carbon burial (Bulian et al., 2022;Bulian et al., 2021;Palcu et al., 2021). Understanding the effect of the progressive restriction already before the onset of evaporite deposition through a detailed multiple sulfur isotopic study of pre-MSC sediments would allow to understand how restriction, that induced a salinity rise and the upward migration of the depth of oxygen depletion, affected the redox cycling of sulfur in Mediterranean basins, resulting in an evolution from : (i) normal marine sediment to (ii) sediment hosting fauna indicative of dysoxic bottom water, (iii) the accumulation of sapropel deposit, and finally (iv) the deposition of evaporites in Mediterranean basins (Bulian et al., 2022). -Aloisi, G., Guibourdenche, L., Natalicchio, M., [START_REF] Aloisi | The geochemical riddle of "low-salinity gypsum" deposits[END_REF] The geochemical riddle of "low-salinity gypsum" deposits. Geochimica et Cosmochimica Acta.

-Alt, J.C. (1995) Sulfur isotopic profile through the oceanic crust: Sulfur mobility and seawater-crustal sulfur exchange during hydrothermal alteration. Geology 23, 585-588.

-Andreetto, F., Aloisi, G., Raad, F., Heida, H., Flecker, R., Agiadi, K., Lofi, J., Blondel, S., Bulian, F., Camerlenghi, A., Caruso, A., Ebner, R., Garcia-Castellanos, D., Gaullier, V., Guibourdenche, L., Gvirtzman, Z., Hoyle, T.M., Meijer, P.T., Moneron, J., Sierro, F.J., Travan, G., Tzevahirtzian, A. -Chen, Y., Zhang, W., Nian, X., Sun, Q., Ge, C., Hutchinson, S.M., Cheng, Q., Wang, F., Chen, J. and Zhao, X. (2021) Greigite as an indicator for salinity and sedimentation rate change: evidence from the Yangtze River delta, China. Journal of Geophysical Research: Solid Earth 126, e2020JB021085.

-Choppala, G., Bush, R., Moon, E., Ward, N., Wang, Z., Bolan, N. and Sullivan, L. ( 2017 -Dunne, J., Sarmiento, J. and Gnanadesikan, A. ( 2007) A synthesis of global particle export from the surface ocean and cycling through the ocean interior and on the seafloor. Global Biogeochemical Cycles -GLOBAL BIOGEOCHEM CYCLE 21.

-Duverger, A., Berg, J.S., Busigny, V., Guyot, F., Bernard, S. and Miot, J. (2020) Mechanisms of Pyrite Formation Promoted by Sulfate-Reducing Bacteria in Pure Culture. Frontiers in Earth Science 8.

-Ebelmen, J.-J. (1845) Sur les produits de la décomposition des espèces minérales de la famille des silicates, Annales des Mines, p. 66.

- -Findlay, A.J., Boyko, V., Pellerin, A., Avetisyan, K., Guo, Q., Yang, X. and -Gennari, R., Manzi, V., Angeletti, L., Bertini, A., Biffi, U., Ceregato, A., Faranda, C., Gliozzi, E., Lugli, S., Menichetti, E., Rosso, A., Roveri, M. and Taviani, M. ( 2013 -Goldhaber, M. and Kaplan, I. (1974) The sulfur cycle. . The Sea Vol. 5, 569-655.

-Gomes, M.L. and Hurtgen, M.T. ( 2013) Sulfur isotope systematics of a euxinic, lowsulfate lake: Evaluating the importance of the reservoir effect in modern and ancient oceans. Geology 41, 663-666.

-Gomes, M.L. and Hurtgen, M.T. (2015) Sulfur isotope fractionation in modern euxinic systems: Implications for paleoenvironmental reconstructions of paired sulfatesulfide isotope records. Geochimica et Cosmochimica Acta 157, 39-55.

-Gomes, M.L., Klatt, J.M., Dick, G.J., Grim, S.L., Rico, K.I., Medina, M., Ziebis, W -Jovovic, I., Grossi, V., Adam, P., Simon, L., Antheaume, I., Gelin, F., Ader, M. and Cartigny, P. (2020) Quantitative and specific recovery of natural organic and mineral sulfur for (multi-)isotope analysis. Organic Geochemistry 146, 104055.

-Kagoshima, T., Sano, Y., Takahata, N., Maruoka, T., Fischer, T.P. and Hattori, K. ( 2015 -Keely, B.J., Blake, S.R., Schaeffer, P. and Maxwell, J.R. (1995) Distributions of pigments in the organic matter of marls from the Vena del Gesso evaporitic sequence. Organic Geochemistry 23, 527-539.

-Kenig, F., Damsté, J.S., Frewin, N.L., Hayes, J.M. and De Leeuw, J.W. (1995) Molecular indicators for palaeoenvironmental change in a Messinian evaporitic sequence (Vena del Gesso, Italy). II: High-resolution variations in abundances and 13C contents of free and sulphur-bound carbon skeletons in a single marl bed. Org Geochem 23, 485-526.

-Knittel, K. and Boetius, A. ( 2009) Anaerobic oxidation of methane: progress with an unknown process. Annu Rev Microbiol 63, 311-334.

-Kölling, M., Bouimetarhan, I., Bowles, M.W., Felis, T., Goldhammer, T., Hinrichs, K.-U., Schulz, M. and Zabel, M. ( 2019) Consistent CO2 release by pyrite oxidation on continental shelves prior to glacial terminations. Nature Geoscience 12, 929-934.

-Kontakiotis, G., Karakitsios, V., Maravelis, A.G., Zarkogiannis, S.D., Agiadi, K., Antonarakou, A., Pasadakis, N. and Zelilidis, A. ( 2021) Integrated isotopic and organic geochemical constraints on the depositional controls and source rock quality of the Neogene Kalamaki sedimentary successions (Zakynthos Island, Ionian Sea). Mediterranean Geoscience Reviews 3,.

-Kotopoulou, E., Godelitsas, A., Göttlicher, J., Steininger, R., Price, R., Fike, D.A., Amend, J.P., Gilhooly, W.P., Druschell, G., Nomikou, P., Gamaletsos, P.N. and Lozios, S. ( 2022 -Kuroda, J., Jiménez-Espejo, F.J., Nozaki, T., Gennari, R., Lugli, S., Manzi, V., Roveri, M., Flecker, R., Sierro, F.J. and Yoshimura, T. ( 2016) Miocene to Pleistocene osmium isotopic records of the Mediterranean sediments. Paleoceanography 31, 148-166.

-Labidi, J., Cartigny, P., Birck, J.L., Assayag, N. and Bourrand, J.J. ( 2012 -Liu, J., Antler, G., Pellerin, A., Izon, G., Dohrmann, I., Findlay, A.J., Røy, H., Ono, S., Turchyn, A.V., Kasten, S. and Jørgensen, B.B. (2021a) Isotopically "heavy" pyrite in marine sediments due to high sedimentation rates and non-steady-state deposition.

Geology.

-Liu, J., Antler, G., Pellerin, A., Izon, G., Dohrmann, I., Findlay, A.J., Røy, H., Ono, S., -Martín, J., Braga, J., Aguirre, J. and Puga-Bernabéu, Á. ( 2009) History and evolution of the North-Betic Strait (Prebetic Zone, Betic Cordillera): A narrow, early Tortonian, tidaldominated, Atlantic-Mediterranean marine passage. Sedimentary Geology 216, 80-90.

-Martín, J., Braga, J. and Betzler, C. (2001) The Messinian Guadalhorce corridor: The last northern, Atlantic-Mediterranean gateway. Terra Nova 13, 418-424.

-Martín, J., Puga-Bernabéu, Á., Aguirre, J. and Braga, J. ( 2014 -Middelburg, J.J., Soetaert, K. and Herman, P.M.J. -Pasquier, V., Sansjofre, P., Rabineau, M., Revillon, S., Houghton, J. and Fike, D.A. ( 2017 -Sabino, M., Dela Pierre, F., Natalicchio, M., Birgel, D., Gier, S. and Peckmann, J. ( 2021)

The response of water column and sedimentary environments to the advent of the Messinian salinity crisis: insights from an onshore deep-water section (Govone, NW Italy). Geological Magazine.

- -Sakai, H. (1971) Sulfur and oxygen isotopic study of barite concretions from banks in the Japan Sea off the Northeast Honshu, Japan. GEOCHEMICAL JOURNAL 5, 79-93.

- -Thullner, M., Dale, A. and Regnier, P. (2009) Global-scale quantification of mineralization pathways in marine sediments: A reaction-transport modeling approach. Geochemistry Geophysics Geosystems 10, Q10012.

- -Weinheimer, h. (1978) stable isotopes of oxygen and carbon in carbonates and organic material from pleistocene to upper miocene sediments at site 374 (dsdp leg 42a).

- -Zúñiga, D., García-Orellana, J., Calafat, A., Price, N.B., Adatte, T., Sanchez-Vidal, A., Canals, M., Sanchez-Cabeza, J.A., Masqué, P. and Fabres, J. ( 2007) Late Holocene finegrained sediments of the Balearic Abyssal Plain, Western Mediterranean Sea. Marine Geology 237, 25-36. Gypsum (CaSO 4 Á2H 2 O) accounts for approximately 20% of the minerals in giant salt deposits formed via the evaporation of seawater throughout Earth's history [START_REF] Zharkov | History of Paleozoic Salt Accumulation[END_REF][START_REF] Warren | Evaporites, Sediments, Resources and Hydrocarbons[END_REF]. Its ionic components SO 4

2À and Ca 2+ are the fourth and fifth most abundant dissolved ions of seawater [START_REF] Stumm | Aquatic Chemistry: Chemical Equilibria and Rates in Natural Waters[END_REF]. By generating fluxes of sulfate and calcium to and from the ocean, continental gypsum weathering and marine gypsum deposition contribute in regulating ocean chemistry [START_REF] Spencer | Control of Seawater composition by mixing of river waters and mid-ocean ridge hydrothermal brines Special Publication[END_REF][START_REF] Hardie | Secular variation in seawater chemistry: An explanation for the coupled secular variation in the mineralogies of marine limestones and potash evaporites over the past 600 m[END_REF][START_REF] Hansen | Cretaceous and Cenozoic evolution of seawater composition, atmospheric O 2 and CO 2 :a model perspective[END_REF][START_REF] Wortmann | Rapid variability of seawater chemistry over the past 130 million years[END_REF][START_REF] Prince | Evaporite deposition in the mid-Neoproterozoic as a driver for changes in seawater chemistry and the biogeochemical cycle of sulfur[END_REF][START_REF] Turchyn | Seawater Chemistry Through Phanerozoic Time[END_REF], as well as atmosphere oxygenation, atmospheric pCO 2 and climate (Wortmann and Chernyavsky, 2007;[START_REF] Wortmann | Rapid variability of seawater chemistry over the past 130 million years[END_REF]Halevy et al., 2012;Shields and Mills, 2020) on geological time scales. The study of gypsum formation in salt giants is therefore relevant to the understanding of the long-term geochemical evolution of Earth's surface.

In the sequence of minerals formed by evaporation of seawater that is often cited in textbooks [START_REF] Rouchy | Les Evaporites -Mate ´riaux singuliers, milieux extre ˆmes[END_REF][START_REF] Warren | Evaporites, Sediments, Resources and Hydrocarbons[END_REF]; Ba z bel and Schreiber, 2014), gypsum is the second precipitate, after carbonates and before halite. Seawater evaporation experiments [START_REF] Usiglio | Analyse de l'eau de la Mediterrane ´e sur les co ˆtes de France[END_REF][START_REF] Valyashko | Playa lakes -a necessary stage in the development of a salt-bearing basin[END_REF][START_REF] Herrmann | Geochemistry of modern seawater and brines from salt pans: main components and bromine distributions[END_REF] and thermodynamic solubility calculations [START_REF] Van't Hoff | Untersuchungen u ¨ber die Bildungsverha ¨ltnisse der ozeanischen Salzblagerungen[END_REF][START_REF] Stewart | Marine Evaporites[END_REF][START_REF] Braitsch | Salt Deposits, Their Origin and Composition[END_REF]Harvie and Weare, 1980;Harvie et al., 1980) show that, in this ''classical" scenario where gypsum saturation is attained by evaporative extraction of fresh water, the point of gypsum saturation is reached when the salinity has risen to about three times the salinity of modern ocean water ($110 g kg À1 ). This salinity threshold is the same in: (i) closed systems where the only water flux is evaporative loss; (ii) constantvolume systems where the water lost by evaporation is replenished by influx from the ocean; and (iii) open systems with leakage, where, in addition to evaporative loss and influx from the ocean, brine is lost to the ocean via a density-driven counter flux [START_REF] Sanford | Brine evolution and mineral deposition in hydrologically open evaporite basins[END_REF]. In summary, in evaporitic systems where the evaporating fluid is seawater (marine evaporites) [START_REF] Hardie | Evaporites: Marine or Non-Marine[END_REF], gypsum deposition is expected to start at a salinity of $110 g kg À1 and proceed until the beginning of the precipitation of halite ($320 g kg À1 ) ( Harvie et al., 1980).

Geochemical evidence from gypsum contained in the Mediterranean Salt Giant suggests that this scenario is not always true. The Mediterranean Salt Giant is Earth's most recent giant salt deposit (Ryan, 2009;Roveri et al., 2014); it formed between 5.97 and 5.33 Ma at the end of the Messinian age of the Miocene epoch, when the Mediterranean Sea underwent an extreme environmental change known as the Messinian Salinity Crisis [START_REF] Hsu | Late Miocene Dessiccation of the Mediterranean[END_REF]Ryan, 2009;Roveri et al., 2014). During this event, a huge salinity increase led to the deposition of a salt-dominated sedimentary body that occupies a volume of 1 million km 3 (Haq et al., 2020). Mediterranean Salt Giant-related deposits are found both in the offshore deep Mediterranean basin (Haq et al., 2020) and in onshore intermediate-toshallow marginal Mediterranean basins (Lugli et al., 2010;Roveri et al., 2014)( Fig. 1). The geochemical signature of gypsum deposited in some marginal basins [START_REF] Natalicchio | Did Late Miocene (Messinian) gypsum precipitate from evaporated marine brines? Insights from the Piedmont Basin (Italy)[END_REF][START_REF] Evans | Coupled measurements of delta O-18 and delta D of hydration water and salinity of fluid inclusions in gypsum from the Messinian Yesares Member, Sorbas Basin (SE Spain)[END_REF][START_REF] Costanzo | Messinian twinned selenite from the Catanzaro Trough, Calabria, Southern Italy: field, petrographic and fluid inclusion perspectives[END_REF] suggest its formation from water masses with a salinity much lower than that ($110 g kg À1 ) predicted by evaporation experiments [START_REF] Usiglio | Analyse de l'eau de la Mediterrane ´e sur les co ˆtes de France[END_REF][START_REF] Valyashko | Playa lakes -a necessary stage in the development of a salt-bearing basin[END_REF][START_REF] Herrmann | Geochemistry of modern seawater and brines from salt pans: main components and bromine distributions[END_REF] and thermodynamic models ( [START_REF] Van't Hoff | Untersuchungen u ¨ber die Bildungsverha ¨ltnisse der ozeanischen Salzblagerungen[END_REF][START_REF] Stewart | Marine Evaporites[END_REF][START_REF] Braitsch | Salt Deposits, Their Origin and Composition[END_REF]Harvie and Weare, 1980;Harvie et al., 1980). In this paper, we will refer to this geochemically peculiar gypsum as ''low-salinity gypsum".

Evidence for low-salinity gypsum in the Messinian Salinity Crisis comes primarily from the study of gypsum primary fluid inclusions. These suggest average salinities of gypsum parent waters equal to 16 g kg À1 in the Piedmont Basin (northern Italy) [START_REF] Natalicchio | Did Late Miocene (Messinian) gypsum precipitate from evaporated marine brines? Insights from the Piedmont Basin (Italy)[END_REF], 32 g kg À1 in the Sorbas basin (Spain) [START_REF] Evans | Coupled measurements of delta O-18 and delta D of hydration water and salinity of fluid inclusions in gypsum from the Messinian Yesares Member, Sorbas Basin (SE Spain)[END_REF] and 27 g kg À1 in in the Catanzaro Trough (southern Italy) [START_REF] Costanzo | Messinian twinned selenite from the Catanzaro Trough, Calabria, Southern Italy: field, petrographic and fluid inclusion perspectives[END_REF]. These values are between 14% and 29% of the threshold salinity at which gypsum precipitates from evaporating seawater in the ''classical" scenario cited above. The absence of evidence for recrystallisation of the primary gypsum deposits at both macro-and microscopic scale [START_REF] Natalicchio | Did Late Miocene (Messinian) gypsum precipitate from evaporated marine brines? Insights from the Piedmont Basin (Italy)[END_REF][START_REF] Costanzo | Messinian twinned selenite from the Catanzaro Trough, Calabria, Southern Italy: field, petrographic and fluid inclusion perspectives[END_REF] suggests that the low salinity signature is pristine and not the result of a diagenetic overprint. These data question the purely evaporitic origin of the Mediterranean gypsum deposits: if not exclusively evaporation of seawater, what process concentrated -or contributed to concentratedissolved Ca 2+ and SO 4 2À to the point of gypsum saturation? Deviations from the purely marine evaporitic mineral sequence can take place if the source of evaporating fluid is different from seawater [START_REF] Hardie | Evaporites: Marine or Non-Marine[END_REF]. In the case of the Messinian Salinity Crisis, there is ample geochemical evidence from Sr-isotopes that continental runoff contributed significant volumes of freshwater to the Mediterranean Sea during gypsum deposition (Topper et al., 2011). Supported by these observations, two hypotheses point to high dissolved Ca 2+ and SO 4 2À in a fresh or brackish water source as the trigger for the formation of lowsalinity gypsum during the Messinian Salinity Crisis. The first hypothesis involves the leaching of exposed penecontemporaneous Messinian gypsum [START_REF] Natalicchio | Did Late Miocene (Messinian) gypsum precipitate from evaporated marine brines? Insights from the Piedmont Basin (Italy)[END_REF][START_REF] Evans | Coupled measurements of delta O-18 and delta D of hydration water and salinity of fluid inclusions in gypsum from the Messinian Yesares Member, Sorbas Basin (SE Spain)[END_REF] -with the possible addition of Ca 2+ and SO 4 2À from carbonate and pyrite weathering [START_REF] Berner | Global Environment: Water, Air and Geochemical Cycles[END_REF] -and the re-precipitation in an adjacent marginal basin. The second hypothesis involves input into the Mediterranean of water via the Paratethys (Grothe et al., 2020); this idea relies on the salinity and chemical composition of the modern Caspian Sea that is brackish (salinity $12 g kg À1 compared to a seawater salinity $35 g kg À1 ) but has dissolved Ca 2+ ($8 mmol kg À1 ) and SO 4 2À ($31 mmol kg À1 ) concentrations comparable to those of the modern ocean [START_REF] Clauer | Fluctuations of Caspian Sea level: Beyond climatic variations?[END_REF]. Both these hypotheses remain unevaluated.

Here we present new geochemical data in support of the formation of Mediterranean marginal gypsum deposits from low-salinity (<110 g kg À1 ) water masses. We expand the observations of [START_REF] Natalicchio | Did Late Miocene (Messinian) gypsum precipitate from evaporated marine brines? Insights from the Piedmont Basin (Italy)[END_REF], [START_REF] Evans | Coupled measurements of delta O-18 and delta D of hydration water and salinity of fluid inclusions in gypsum from the Messinian Yesares Member, Sorbas Basin (SE Spain)[END_REF] and [START_REF] Costanzo | Messinian twinned selenite from the Catanzaro Trough, Calabria, Southern Italy: field, petrographic and fluid inclusion perspectives[END_REF] to include novel fluid inclusion salinity data from primary gypsum deposits of the Vena del Gesso basin (northern Italy) and the isotope composition of primary gypsum hydration water (d 18 O H2O and dD H2O ) from the Caltanissetta basin (Sicily), the Vena del Gesso basin and the Piedmont basin (northern Italy), as well as the hydration water isotope (d 18 O H2O and dD H2O ) composition from secondary gypsum from the Volterra basin. Gypsum crystallization water isotopes are complementary to fluid inclusion salinity measurements, since gypsum crystallization water pervades the whole crystalline lattice, rather than being concentrated in spatially distinct portions of the crystal matrix. In addition, we present the isotope composition of sulfur (d 34 S SO4 ) and oxygen (d 18 O SO4 ) of the sulfate ion in gypsum from the Caltanissetta, the Vena del Gesso and Piedmont basins, that provide information on sulfate sources and the biogeochemical sulfur cycle.

We apply a simple geochemical model that simulates the evolution of salinity, gypsum saturation state and water isotope composition (d 18 O and dD) during the evaporation of mixtures of seawater and a Ca 2+ and/or SO 4 2À -bearing freshwater source. The model is used to (i) evaluate the potential of gypsum hydration water isotopes as indicators of the formation of gypsum from low-salinity parent waters, and (ii) evaluate scenarios for low-salinity gypsum formation that involve a Ca 2+ and SO 4 2 -rich fresh water source. Finally, we discuss a complementary scenario for the formation of low-salinity gypsum that involves the spatially-restricted and temporally-transient increase of Ca 2+ and SO 4 2À at the sediment-water interface of the basin triggered by the biogeochemical sulfur cycle.

GEOLOGICAL SETTING, STUDY AREAS AND GYPSUM LITHOFACIES

Geological setting of Mediterranean Salt Giant deposits

The bulk of the evaporitic deposits that make up the Mediterranean Salt Giant (primarily halite and gypsum, with subordinate K-Mg salts such as kainite, carnallite and bishofite) were deposited in the deep Mediterranean basin where they are buried under several hundreds of meters of Plio-Quaternary pelagic sediments (Roveri et al., 2014;Haq et al., 2020). Evaporites (mainly gypsum, with minor halite and K-Mg salts) were also deposited in intermediate-(1000 to 200 m) to shallow-(<200 m) depth marginal basins, now mostly exposed on land (Lugli et al., 2010;Dela Pierre et al., 2014)(Fig. 1). Marginal basin deposits have been extensively studied (Krijgsman et al., 1999[START_REF] Krijgsman | The onset of the Messinian salinity crisis in the Eastern Mediterranean (Pissouri Basin, Cyprus)[END_REF]Rouchy and Caruso, 2006;Lugli et al., 2010;Dela Pierre et al., 2011;[START_REF] Manzi | High-frequency cyclicity in the Mediterrnean Messinian Evaporites: evidence for solar-lunar climate forcing[END_REF]Manzi et al., , 2013Manzi et al., , 2016;;Roveri et al., 2014Roveri et al., , 2019;;Cosentino et al., 2018). They form gypsum-dominated sedimentary bodies that can be followed for tens of kilometers both onshore and in the marginal offshore domain (Driussi et al., 2015;Ochoa et al., 2015;[START_REF] Raad | The Messinian Salinity Crisis deposits in the Balearic Promontory: An undeformed analog of the MSC Sicilian basins? Mar[END_REF]Manzi et al., 2020).

Astronomical dating of marginal onshore successions across the Mediterranean has shown that the onset of the Messinian Salinity Crisis is a synchronous event at 5.97 Ma (Krijgsman et al., 1999;Manzi et al., 2013)o r slightly diachronous (Rouchy and Caruso, 2006;Caruso et al., 2015). Messinian evaporites were originally subdivided into two informal stratigraphic units: Lower Evaporites, including gypsum and the thick massive halite and K-Mg salts of Sicily and deep basins, and Upper Evaporites, mostly consisting of gypsum (Rouchy and Caruso, 2006 and references therein). More recently a different terminology was proposed following the three-stage model of the Messinian Salinity Crisis events that is based on outcrop studies of key section from different Mediterranean sub-basin [START_REF] Ciesm | The Messinian Salinity Crisis from mega deposits to microbiology[END_REF]Roveri et al., 2014). According to this model, whose validity in the entire Mediterranean basin is debated (Meilijson et al., 2019;[START_REF] Artiaga | The Messinian evaporites of the Mesaoria basin (North Cyprus): A discrepancy with the current chronostratigraphic understanding[END_REF], up to 16 precession-controlled shale-gypsum cycles, known as the Primary Lower Gypsum, were deposited in silled marginal basins during the first phase of the Messinian Salinity Crisis . During the second phase of the Messinian Salinity Crisis (5.60-5.55 Ma), the Primary Lower Gypsum deposits were uplifted and partially eroded, resulting in clastic gypsum deposits (Resedimented Lower Gypsum unit) that locally interfinger (Caltanissetta basin) with halite and K-Mg salts (Manzi et al., 2021). Primary gypsum deposition took place again in the southern (Caltanissetta basin) and eastern (Cyprus, Crete) Mediterranean marginal basins in the third phase of the Messinian Salinity Crisis (5.55-5.33 Ma) forming the Upper Gypsum unit (corresponding to the Upper Evaporites) that consist of up to 8 marl/gypsum cycles (Roveri et al., 2014). The third phase of the Messinian Salinity Crisis also contains a non-evaporitic, brackish water deposit called Lago-Mare (Andreetto et al., 2021) that preceded the reconnection of the Mediterranean with the Atlantic Ocean at the Miocene-Pliocene boundary (5.33 Ma).

Study areas

We investigated onshore Primary Lower Gypsum and Upper Gypsum deposits from the Caltanissetta basin: Santa Elisabetta section (37°26 0 50 00 N, 13°32 0 58 00 E; 14 samples), Gibliscemi section (37°12 0 21 00 N, 14°16 0 16 00 E; 6 samples) and Eraclea Minoa section (37°23 0 37 00 N, 13°17 0 28 00 E; 32 samples); Primary Lower Gypsum deposits from the Vena del Gesso basin: Monticino section (44°13 0 29 00 N, 11°4 5 0 43 00 E; 40 samples) and Monte Tondo quarry section (44°15 0 04 00 N, 11°40 0 13 00 E; 5 samples); Primary Lower Gypsum deposits from the Piedmont basin: Banengo section (45°04 0 18.5 00 N8 °03 0 25.2 00 E; 45 samples), Arnulfi section (43°37 0 06 00 N; 7°53 0 41 00 E; 27 samples) and Pollenzo section (44°41 0 08 00 N; 7°55 0 33 00 E; 71 samples) (Fig. 1b). Next to these primary gypsum precipitates, we analyzed Primary Lower Gypsum deposits from the Castellina Quarry (45°25 0 14 0 ' N; 10°32 0 55 0 'E; 6 samples) of the Volterra basin (Tuscany, Italy) that are evidently secondary deposits based on their macroscopic features. A description of the above-mentioned marginal basins and of the investigated lithological sections is provided in the Supplementary Material section. The investigated stratigraphic sections are presented in Fig. 1b.

Gypsum lithofacies

The Primary Lower Gypsum and Upper Gypsum units comprise different gypsum lithofacies that are described below. The names of lithofacies presented in italics in this section are used to describe the samples included in this study and listed in the sample tables S1-S8.

Selenite gypsum

In the Vena del Gesso and Piedmont basins, the lower beds of the Primary Lower Gypsum unit are up to 30 m thick and consist of vertically-oriented, twinned (arrow head and swallow tail), selenite crystals [START_REF] Orti | Selenite facies in marine evaporites: A review[END_REF] nucleated and grown on the basin floor (hereafter, bottom-grown); they form the massive and banded selenite lithofacies. The massive selenite is composed of dm-sized crystals with uniform size throughout the beds. Crystals up to 2.5 m high (giant selenite) are locally observed; the banded selenite is composed of cm-sized palisades if twinned crystals that are separated by mm-thick clay and carbonate layers. In both the massive and the banded selenite lithofacies, the re-entrant angle of the crystals shows an internal lamination produced by the alternation of turbid laminae, rich in solid inclusions (mostly filamentous microfossils and clay-rich floccules), and limpid laminae in which solid inclusions are scarce or absent. Such lamination reflects the influence of short-term (annual, seasonal) climate variability [START_REF] Orti | Selenite facies in marine evaporites: A review[END_REF]Dela Pierre et al., 2015;Reghizzi et al., 2018). The 6th gypsum bed and the overlying beds are characterized by the presence of the branching selenite lithofacies, composed of cm-sized horizontally-oriented crystals grouped together to form asymmetrical conical structures with one side (branch) more developed than the other (Lugli et al., 2010;[START_REF] Natalicchio | Gypsum lithofacies and stratigraphic architecture of a Messinian marginal basin (Piedmont Basin, NW Italy)[END_REF]. In the Vena del Gesso basin, branching selenite, also bottom-grown, is found in the upper part of the gypsum beds, above the massive and banded selenite lithofacies (Lugli et al., 2010). The massive selenite lithofacies is present also in the Caltanisetta basin where it forms the bulk of gypsum deposits in the Santa Elisabetta section (Primary Lower Gypsum), the upper part of the Gibliscemi section and the Eraclea Minoa section (Upper Gypsum).

Laminar gypsum

This latter lithofacies, formerly termed balatino [START_REF] Ogniben | Petrografia della serie solfifera siciliana e considerazioni geologiche[END_REF], consists of the alternation of gypsum rich laminae (1-2 mm thick) and clay and dolomite-rich laminae (100-200 lm thick). The gypsum-rich laminae are composed of interlocked prismatic and acicular crystals ranging in size from a few lm to some hundreds of lm. The laminar gypsum represents an in-situ cumulate deposit, formed by the nucleation of tiny gypsum crystal in the water column and their subsequent deposition on the seafloor [START_REF] Natalicchio | Gypsum lithofacies and stratigraphic architecture of a Messinian marginal basin (Piedmont Basin, NW Italy)[END_REF]. In the Piedmont basins, the laminar gypsum is associated to the branching selenite facies [START_REF] Natalicchio | Gypsum lithofacies and stratigraphic architecture of a Messinian marginal basin (Piedmont Basin, NW Italy)[END_REF]. In sections from the Caltanissetta basin considered in this work, laminar gypsum occurs in the Gibliscemi section and in the Eraclea Minoa section (Upper Gypsum).

Secondary gypsum lithofacies

Following episodes of dehydration-hydration, gypsum is transformed into anhydrite, and anhydrite is subsequently transformed into gypsum (see supplementary material), with a complete loss of the original gypsum lithofacies. The neo-formed gypsum lithofacies include: alabastrine gypsum (an aggregate of gypsum microcrystals), fibrous gypsum (fracture-filling fibrous gypsum crystals) and chicken-wire gypsum (microcrystalline gypsum nodules separated by thin seams of host sediments). In our study, these secondary gypsum lithofacies occur only in the Volterra basin [START_REF] Testa | Gypsum-anhydrite transformations in Messinian evaporites of central Tuscany (Italy)[END_REF].

METHODS

In this section we present a summary of the analytical methods used. A full description of the methods is provided in the Supplementary Materials section.

Sampling

Samples were obtained from the outcrop by mechanically breaking gypsum fragments with a geological hammer. Care was taken to remove a few centimeters of the gypsum surface to avoid sampling the weathered surface layer. A few core samples from the Santa Elisabetta and Gibliscemi sections were obtained using a water-cooled drill with a core-bit diameter of 3 cm. This sampling technique was not used further because it is time-consuming, has the potential of heating the gypsum and did not result in gypsum samples that were visually more pristine than those obtained with the geological hammer. Sample tables are presented in the Supplementary Materials section. Sampling positions appear in Fig. 1b.

Isotopic composition of gypsum hydration water

Gypsum hydration water was extracted following the protocol of [START_REF] Fontes | Fractionnement isotopique dans l'eau de cristallisation du sulfate de calcium[END_REF] (see supplementary methods). Liquid water samples obtained by dehydrating gypsum were analyzed for oxygen (d 18 O) and hydrogen (dD) isotope composition using a Picarro L2130-i (LOCEAN laboratory, Sorbonne Universite ´) or a Picarro L2140-i (IPGP, Universite ´de Paris) laser spectrometer. The isotope composition of gypsum hydration water is expressed in delta notation (d 18 O for oxygen 18 and dD for deuterium) defined as the per mil deviation of the 18 O/ 16 O and D/H ratios in the seawater from the ratio in the Vienna Standard Mean Ocean Water (VSMOW). The standard deviation of the isotope values of multiple analyses of the same gypsum hydration water sample are typically smaller than 0.05‰ for d 18 O and 0.2‰ for dD.

Isotopic composition of the sulfate ion of gypsum

Gypsum samples were ground and dissolved in deionized and acidified water. The solution was then filtered (0.8 lm), heated and the dissolved sulfate was precipitated as barium sulfate by addition of a 1 N-BaCl 2 solution. The precipitated BaSO 4 was filtered (0.8 lm), washed with deionized water and dried in an oven at 90 °C.

The d 34 S and d 18 O SO4 of gypsum sulfate from the Piedmont and Caltanissetta basins were measured at the LOCEAN laboratory, Sorbonne Universite ´. Isotope analyses were carried out in triplicate on a Thermo Scientific Delta V Advantage IRMS on CO (d 18 O) and SO 2 (d 34 S) produced respectively by pyrolysis (using a Conflo IV interface, on 150 lg of barium sulphate plus 1 mg of AgCl) and combustion (using a Flash-EA2000 Thermal Conversion Elemental Analyzer unit, on 250 lg of barium sulfate plus 1mgofV 2 O 5 ) of BaSO 4 obtained by chemical precipitation of sulfate gypsum in a BaCl solution. The d 18 O and d 34 S values were calibrated using the IAEA international reference NBS 127 (d 18 O = +8.59 ± 0.2‰ vs SMOW and d 34 S = +21.12 ± 0.22‰ vs CDT [START_REF] Brand | Assessment of international reference materials for isotope-ratio analysis (IUPAC Technical Report)[END_REF].

The d 18 O SO4 of gypsum sulfate from the Vena del Gesso basin was measured at the Universite ´de Bourgogne (Dijon) on purified barite samples obtained with the procedure described above. Isotope measurements were carried out in duplicate on an Elementar vario PYRO cube elemental analyzer in-line with an Isoprime 100 mass spectrometer in continuous flow mode. Analytical errors are ±0.4‰ (2r) based on replicate analyses (n = 21) of the international barite standard NBS-127, which was used for data correction via standard-sample-standard bracketing.

d 34 S SO4 of gypsum sulfate from the Vena del Gesso basin was measured at IPGP (Paris) on gaseous SF 6 with a MAT-253 dual inlet mass spectrometer. The gypsum powder was first reduced with a ''Thode" solution into H 2 S [START_REF] Thode | Sulphur Isotope Geochemistry[END_REF] that was subsequently trapped in NaOH solution and then precipitated into Ag 2 S by adding a AgNO 3 solution [START_REF] Geng | A simple and reliable method reducing sulfate to sulfide for multiple sulfur isotope analysis[END_REF]. The entire extraction procedure gave sulfur yields between 76% and 116% (n = 29) with a mean of 95 ± 2% (1r), with the lowest yields explained by higher amount of non-sulfur bearing mineral phases such as clays. Ag 2 S was then introduced into Ni reaction bombs, reacted with F 2 at 310 °C overnight to produce SF 6 and then purified cryogenically and by gas chromatography (Ono et al., 2006;[START_REF] Labidi | Determination of multiple sulfur isotopes in glasses: A reappraisal of the MORB delta S-34[END_REF]. The SF 6 was then introduced into the MAT-253 for isotopic measurements that yielded analytical precisions of 0.2‰ (accounting for full reproducibility of the steps described above). Sulfur and oxygen isotope data are expressed in delta notation defined as the per mil deviation of the 34 S/ 32 S and 18 O/ 16 O ratios from the international CDT and SMOW standards. The analytical precision is ±0.1‰ and ±0.2‰ for d 34 S and d 18 O measurements respectively.

Salinity of gypsum fluid inclusions

Primary fluid inclusions were studied from 6 bottomgrown gypsum samples from the Monte Tondo section (Vena del Gesso basin). A total of 60 microthemometric measurements were carried out using a calibrated Linkam THMSG600 heating-freezing stage attached to an Olympus polarizing microscope using the method described in [START_REF] Natalicchio | Did Late Miocene (Messinian) gypsum precipitate from evaporated marine brines? Insights from the Piedmont Basin (Italy)[END_REF] and explained in the Supplementary Material section. For each analyzed sample, several micro-samples (fragments) from the re-entrant angle of gypsum twins were investigated for their fluid inclusions. Salinities of the fluid inclusions were calculated from the final melting temperatures of ice and expressed as weight % NaCl equivalent [START_REF] Bodnar | Revised Equation and Table for Determining the Freezing-Point Depression of H 2 O-Nacl Solutions[END_REF] and as ppt of seawater [START_REF] Goldstein | Systematics of Fluid Inclusions in Diagenetic Minerals[END_REF]. Petrographic and cathodoluminescence observations were carried out by plane-polarized and cross-polarized light microscopy using a CITL 8200 mk3 equipment, operating at about 17 kV and 400 lA.

Mixing-evaporation model

We used a numerical model to simulate the evolution of salinity and isotope composition (d 18 O H2O , dD H2O , 87 Sr/ 86 Sr and d 34 S SO4 ) of mixtures of Atlantic Ocean water and freshwater from continental runoff subject to evaporation until the point of gypsum saturation (X Gypsum = 1). The model is a simple mixing-evaporation model in which the mixing and evaporation processes are considered successively, rather than contemporaneously, and where the compositional change in time is produced exclusively by the decrease of the total amount of water in the system. The continental runoff endmember can be ion-free or can contain dissolved Ca 2+ and SO 4 2À , reflecting the addition of these ions via gypsum, carbonate and pyrite weathering. The model considers initial fluid mixtures that have dissolved Ca 2+ /Cl À and SO 4 2À /Cl À ratios ranging from those of seawater (0.0188 and 0.052, respectively) to ratios ten times larger, while the other ion/Cl À ratios are held equal to those of seawater. The model calculates the saturation state with respect to gypsum as a function of water temperature (5-35 °C), ionic strength and Ca 2+ and SO 4 2À concentration by using an interpolation function based on the results of Pitzer ion-interaction calculations that include all major ions of seawater and are designed to cover all the chemical compositions explored with the evaporation model. Although carbonate alkalinity (dissolved HCO 3 and CO 3 2-) is included in the Pitzer system, to simplify the evaporation model we do not consider precipitation of carbonate minerals in our simulations. This simplification has only small effects on model results since only $1 mM of dissolved Ca 2+ would be sequestered by seawater dissolved total alkalinity ($2 mM) prior to reaching gypsum saturation. The isotope composition (d 18 O H2O , dD H2O , 87 Sr/ 86 Sr and d 34 S SO4 ) of gypsum samples from the Primary Lower Gypsum and Upper Gypsum units discussed in the present work is used, in conjunction with the model, to evaluate existing scenarios for the formation of low-salinity gypsum deposits. The model is compiled in the software Wolfram Mathematica 12 and is run on an Apple iMac desktop computer. A full description of the mixing-evaporation model and the Pitzer model are provided in the supplementary material section.

ANALYTICAL RESULTS

A total of 60 new fluid inclusion salinity measurements were obtained from 6 samples of vertically oriented selenite crystals from the Primary Lower Gypsum unit of the Vena del Gesso basin (Primary Lower Gypsum cycles [START_REF] Lein | Isotopic composition of sulfur compounds in sediments of the Pacific[END_REF]4,6,10,[START_REF] Aloisi | The geochemical riddle of ''low-salinity gypsum" deposits[END_REF] (Supplementary Table 9). The salinity of single fluid inclusion varies from 5.5 to 179 g kg À1 (Fig. 2a). The large range of salinities measured within the same sample is related to primary fluid inclusions belonging to different growth bands, likely suggesting salinity fluctuations during gypsum growth. The average salinity of the 6 samples varies from 24.5 to 78.2 g kg À1 , while the average of all salinity measurements is equal to 52.2 g kg À1 . Compared to the average fluid inclusion salinity data from Primary Lower Gypsum unit of the Piedmont basin (16 g kg À1 , [START_REF] Natalicchio | Did Late Miocene (Messinian) gypsum precipitate from evaporated marine brines? Insights from the Piedmont Basin (Italy)[END_REF], the Upper Gypsum unit of the Catanzaro Trough (27 g kg À1 ; [START_REF] Costanzo | Messinian twinned selenite from the Catanzaro Trough, Calabria, Southern Italy: field, petrographic and fluid inclusion perspectives[END_REF] and the Primary Lower Gypsum unit of the Sorbas basin (32.3 g kg À1 ; [START_REF] Evans | Coupled measurements of delta O-18 and delta D of hydration water and salinity of fluid inclusions in gypsum from the Messinian Yesares Member, Sorbas Basin (SE Spain)[END_REF], the Vena del Gesso basin Primary Lower Gypsum fluid inclusions are more saline (Table 1).

We calculated the isotope composition of gypsum parent waters from the measured isotope composition of gypsum hydration water, using the water-gypsum oxygen and hydrogen fractionation factors determined by Gazquez represent the frequency (left vertical axis) of salinity measurements made in primary fluid inclusion from gypsum from the Piedmont and Vena del Gesso basins (northern Italy) (this work), the Catanzaro Trough (Calabria, Southern Italy) [START_REF] Costanzo | Messinian twinned selenite from the Catanzaro Trough, Calabria, Southern Italy: field, petrographic and fluid inclusion perspectives[END_REF] and the Sorbas basin [START_REF] Evans | Coupled measurements of delta O-18 and delta D of hydration water and salinity of fluid inclusions in gypsum from the Messinian Yesares Member, Sorbas Basin (SE Spain)[END_REF]; blue circles represent the salinity-d 18 O composition of Sorbas basin gypsum fluid inclusions [START_REF] Evans | Coupled measurements of delta O-18 and delta D of hydration water and salinity of fluid inclusions in gypsum from the Messinian Yesares Member, Sorbas Basin (SE Spain)[END_REF]. Coloured circles in panels (b) to (e) are the isotope composition (d 18 O H2O and dD H2O ) of gypsum parent waters from the Caltanissetta (Sicily, Italy), Vena del Gesso (Apennines, Italy), Piedmont (Italy) (this work) and Sorbas (Spain) [START_REF] Evans | Coupled measurements of delta O-18 and delta D of hydration water and salinity of fluid inclusions in gypsum from the Messinian Yesares Member, Sorbas Basin (SE Spain)[END_REF]. Crosses, and the vertical grey band in panel f, indicate the isotope composition of parent waters of gypsum formed in modern salt works of the Mediterranean region [START_REF] Fontes | Fractionnement isotopique dans l'eau de cristallisation du sulfate de calcium[END_REF][START_REF] Fontes | Fractionnement isotopique de l'hydroge `ne dans l'eau de cristallisation du gypse[END_REF]Longinelli, 1979;[START_REF] Evans | Coupled measurements of delta O-18 and delta D of hydration water and salinity of fluid inclusions in gypsum from the Messinian Yesares Member, Sorbas Basin (SE Spain)[END_REF]. The black dot is the isotope composition of surface Atlantic water in from of Gibraltar (Benetti et al., 2017a). The continuous green line, followed by the green dotted line in panels (b) to (f), represent the model evolution in the d 18 O H2O -dD H2O space during the evaporation of Mediterranean Sea water under evaporative conditions based on in-situ atmospheric measurements from the Vaccare `s Lagoon (Camargue area, southern France) [START_REF] Delattre | Deuterium excess in the atmospheric water vapour of a Mediterranean coastal wetland: regional vs. local signatures[END_REF]. The green dot with a black circle separating the continuous green line from the green dotted line represents the gypsum saturation point; the d 18 O H2O -dD H2O evolution after the gypsum saturation point (green dotted line) reflects the increasing importance of fractionation between free water molecules and water molecules within the hydration sphere of dissolved ions [START_REF] Sofer | Isotopic composition of hydration water in gypsum[END_REF]Gat, 1972, 1975). Note: MWL -Meteoric Water Line [START_REF] Craig | Deuterium and oxygen 18 variations in the ocean and the marine atmosphere[END_REF].

G. [START_REF] Fontes | Fractionnement isotopique dans l'eau de cristallisation du sulfate de calcium[END_REF][START_REF] Fontes | Fractionnement isotopique de l'hydroge `ne dans l'eau de cristallisation du gypse[END_REF]Longinelli, 1979;[START_REF] Evans | Coupled measurements of delta O-18 and delta D of hydration water and salinity of fluid inclusions in gypsum from the Messinian Yesares Member, Sorbas Basin (SE Spain)[END_REF], to values 3.6‰ lighter than Atlantic waters in front of Gibraltar (Benetti et al., 2017a). In the Piedmont basin, the gypsum parent waters are markedly lighter (À7. 2e; [START_REF] Evans | Coupled measurements of delta O-18 and delta D of hydration water and salinity of fluid inclusions in gypsum from the Messinian Yesares Member, Sorbas Basin (SE Spain)[END_REF] are intermediate between those of the Vena del Gesso and the Piedmont basins. Except from the two massive selenite samples that have similar isotopic values than gypsum from the Piedmont basin, the parent waters of alabastrine, fibrous and chicken-wire gypsum samples from the Castellina Marittima Quarry in the Volterra basin show the typical deuterium enrichment described by [START_REF] Sofer | Isotopic composition of hydration water in gypsum[END_REF] and [START_REF] Bath | 18 O/ 16 O and 2 H/ 1 H chages during progressive hydration of a Zechstein anhydrite formation[END_REF] for gypsum formed diagenetically by hydration of anhydrite (Fig. 2f). Fig. 3 presents the d 34 S and d 18 O of the dissolved sulfate ion in equilibrium with 127 gypsum samples from the investigated Primary Lower Gypsum and Upper Gypsum units, considering the sulfur (Thode and Monster, 1965;[START_REF] Raab | Sulfur isotopic variations during seawater evaporation with fractional crystallization[END_REF] and oxygen (Lloyd, 1968;[START_REF] Van Driessche | Unraveling the Sulfate Sources of (Giant) Gypsum Crystals Using Gypsum Isotope Fractionation Factors[END_REF] isotope fractionations between the dissolved sulfate ion and gypsum (e 34 S SO4-GYP = À1.65‰ and e 18 O SO4 -GYP = À3.6‰, respectively). Most of the d 34 S values of dissolved SO 4 2À deduced from Primary Lower Gypsum and Upper Gypsum deposits span a range of 2‰, centered around 22‰ vs CDT, which corresponds to the d 34 S of the SO 4 2À ion dissolved in the Messinian ocean (Masterson et al., 2016). The d 18 O compositions of dissolved SO 4 2À deduced from Primary Lower Gypsum and Upper Gypsum deposits span a larger range of about 8‰ (excluding one very 18 O-depleted Primary Lower Gypsum sample from the Arnulfi section), from 6 to 15‰ vs SMOW, which includes the d 18 O of the SO 4 2À ion dissolved in the Messinian Ocean (Turchyn and Schrag, 2004). The dissolved sulfate ion corresponding to Messinian gypsum is isotopically heavier than that corresponding to older gypsum deposits (>14 million years old) formed from evaporation of ocean water that was considerably more 34 S-and 18 O-depleted compared to the Messinian ocean (Utrilla et al., 1992)( Fig. 3). Compared to the range of d 34 S SO4 and d 18 O values in modern sediment pore waters where microbial sulfate reduction is active (blue area in Fig. 3) [START_REF] Fotherby | Modelling the Effects of Non-Steady State Transport Dynamics on the Sulfur and Oxygen Isotope Composition of Sulfate in Sedimentary Pore Fluids[END_REF], the range of d 34 S SO4 corresponding to Primary Lower Gypsum and Upper Gypsum deposits is limited (Fig. 3).

MODELLING: APPROACHES AND RESULTS

Modelling the evaporation of seawater

We first tested the ability of the model to reproduce the isotope composition of gypsum (d 18 O H2O and dD H2O ) formed in modern Mediterranean salt works [START_REF] Fontes | Fractionnement isotopique dans l'eau de cristallisation du sulfate de calcium[END_REF][START_REF] Fontes | Fractionnement isotopique de l'hydroge `ne dans l'eau de cristallisation du gypse[END_REF]Longinelli, 1979;[START_REF] Evans | Coupled measurements of delta O-18 and delta D of hydration water and salinity of fluid inclusions in gypsum from the Messinian Yesares Member, Sorbas Basin (SE Spain)[END_REF]( Fig. 2). In this base-case evaporation scenario, the atmospheric conditions that control isotope effects during evaporation (Supplementary Table 11) are based on an isotopic study of atmospheric water vapor in the Vaccare `s Lagoon (Camargue Area, southern France) [START_REF] Delattre | Deuterium excess in the atmospheric water vapour of a Mediterranean coastal wetland: regional vs. local signatures[END_REF], located 25 km away from the Aigues-Mortes salt works [START_REF] Fontes | Fractionnement isotopique dans l'eau de cristallisation du sulfate de calcium[END_REF][START_REF] Fontes | Fractionnement isotopique de l'hydroge `ne dans l'eau de cristallisation du gypse[END_REF])( Fig. 1). These conditions include an elevated water temperature (29 °C), a low atmospheric relative humidity (0.46) and a low wind speed (1.43 m s À1 ). Neglecting mixing with continental runoff, modern Atlantic Ocean water in front of Gibraltar (d 18 O=1 ‰ vs SMOW; dD=6 ‰ vs SMOW, salinity = 36.4 g/kg; Benetti et al. 2017a) was allowed to evaporate. The path in the d 18 O-dD space is shown as green lines in the panels of Fig. 2, while the gypsum saturation point is marked as a green dot within a black circle at the end of the evaporation path. For this base-case scenario, the gypsum saturation point is attained when salinity is equal to 115 g kg À1 (data not shown), d 18 O H2O is equal to 7.71‰ vs SMOW and dD H2O is equal to 32.32‰ vs SMOW (Fig. 2). This water isotope composition falls within the range of values observed in modern Mediterranean salt Fig. 3. Sulfur (d 34 S) and oxygen (d 18 O) isotope composition of the dissolved sulfate ion deduced from the analyses of Primary Lower Gypsum and Upper Gypsum gypsum deposits taking into account the sulfur (Thode and Monster, 1965) and oxygen [START_REF] Lloyd | Oxygen Isotope Enrichment of Sea Water by Evaporation[END_REF] isotope fractionations between the dissolved sulfate ion and gypsum (e 34 S SO4-GYP = À1.65‰ and e 18 O SO4 -GYP = À3.6‰, respectively). Previously published Primary Lower Gypsum data (Lu et al., 2001;[START_REF] Evans | Coupled measurements of delta O-18 and delta D of hydration water and salinity of fluid inclusions in gypsum from the Messinian Yesares Member, Sorbas Basin (SE Spain)[END_REF][START_REF] Garcia-Veigas | Late Miocene evaporite geochemistry of Lorca and Fortuna basins (Eastern Betics, SE Spain): Evidence of restriction and continentalization[END_REF] has been corrected, where, appropriate, taking into account the latest reported values of d 34 S and d 18 O of the NBS 127 barium sulfate standard [START_REF] Brand | Assessment of international reference materials for isotope-ratio analysis (IUPAC Technical Report)[END_REF]. Data from Mesozoic and Cenozoic gypsum deposits from southern Spain (''ancient gypsum > 14 Ma") was taken from Utrilla et al. (1992).

G. Aloisi et al. / Geochimica et Cosmochimica Acta xxx (2022) works, showing that the model correctly simulates this relatively constrained evaporative system. The two key parameter settings that result in a good fit of the numerical model to the salt work gypsum parent water isotopes are a low atmospheric relative humidity (46%) and a low wind speed (1.4 m/s). Both these parameter settings produce a strong deuterium excess in the evaporative flux and thus a strong deuterium depletion of the residual water body (see Supplementary Material). The evolution of salinity and water isotopes after the point of gypsum saturation, when water isotopes are affected considerably by the fractionation between free water molecules and water molecules in the hydration sphere of dissolved ions [START_REF] Sofer | Isotopic composition of hydration water in gypsum[END_REF]Gat, 1972, 1975), is represented by a green dotted line. Note that the influence of the salt effect becomes important only after the gypsum saturation point is attained.

The model was then forced considering the physical and chemical boundary conditions representative of the modern open Mediterranean Sea that are different from the more continental ones of the Vaccare `s Lagoon (Supplementary Fig. 2). These boundary conditions are taken from the geographically-resolved (0.25 degree intervals for both latitude and longitude), monthly averages of climate variables (sea-surface temperature, air temperature, wind speed, atmospheric pressure and atmospheric humidity) provided by the ERA5 reanalysis of the European Centre for Medium-Range Weather Forecast (ECMWF), and from Mediterranean Sea [START_REF] Gat | Isotope composition of air moisture over the Mediterranean Sea: an index of the air-sea interaction pattern[END_REF] and north Atlantic (Benetti et al., 2017b) atmospheric water vapor isotope measurements, as explained in the Supplementary Material. Altogether, we carried out three sets of 1517 model runs each, representative of the geographically-resolved, 30year (1990-2019) average climatology for the months of January and of July, and of the average yearly climatology for the same 30-year period. In Fig. 2f the water isotope composition calculated by the evaporation model at gypsum saturation for these three groups of boundary conditions are shown as blue (average July), green (average January) and red (average yearly) dots. In all of the considered model runs, the evaporation model calculates a salinity of $115 g kg À1 at the gypsum saturation point (data not shown).

Evaporation of Atlantic Ocean waters under the considered Mediterranean Sea conditions results in a range of water isotope compositions at gypsum saturation, ranging from values approaching those of the initial fluid (d 18 O=1 ‰ vs SMOW; dD=6 ‰ vs SMOW) in summer conditions (Fig. 2f, blue dots), to values (>7‰ vs SMOW) heavier than gypsum parent waters formed in salt works in winter conditions (Fig. 2f, green dots). Under July and average yearly conditions, the model water isotope composition at gypsum saturation (Fig. 2f, red dots) has a d 18 O comparable to that of the gypsum parent waters of the Santa Elisabetta and Eraclea Minoa sections of the Caltanissetta basin, the Monte Tondo quarry section of the Vena del Gesso, the Volterra basin and the Banengo section of the Piedmont basin (Fig. 2; Table 1), although it is up to 20‰ enriched in deuterium compared to gypsum parent waters of these Messinian sections. The calculated water isotope composition under average yearly conditions, however, are considerably heavier that the isotope compositions of gypsum parent waters from the Arnulfi and Pollenzo sections of the Piedmont basin (Fig. 2; Table 1). These results put limits to the use of gypsum parent water isotope compositions as indicators of gypsum parent water salinity (see discussion in Section 6.2).

Modelling the evaporation of mixtures of seawater and continental runoff

Since there is ample geochemical evidence that continental runoff contributed significant volumes of water to the Mediterranean Sea during the deposition of the Primary Lower Gypsum and Upper Gypsum units (Flecker et al., 2002;Topper et al., 2011Topper et al., , 2014;;Roveri et al., 2014;Reghizzi et al., 2018), we tested the effect on model output of the mixing of ion-free meteoric waters with Atlantic waters prior to evaporation. All the model runs in this section were carried out at a temperature of 20 °C (representative of average Mediterranean surface water temperature; Supplementary Fig. S2), to single out the effect of mixing with continental runoff on model results. We a priori considered the average rain water isotope composition of the Almeria region (IAEA, 2005;[START_REF] Evans | Coupled measurements of delta O-18 and delta D of hydration water and salinity of fluid inclusions in gypsum from the Messinian Yesares Member, Sorbas Basin (SE Spain)[END_REF], where the Sorbas basin is located, mixed it in four different proportions (f R = 0.2, 0.4, 0.6 and 0.9) with Atlantic waters, and calculated the salinity-d 18 O H2O evolution to the point of gypsum saturation (red lines in Fig. 4). The physical conditions forcing evaporation were calculated as the average of all 1517 points used in the 30-year yearly average climatology (Supplementary Table 11). The addition of ion-free meteoric waters results in an initial fluid with a higher degree of under-saturation with respect to gypsum, and thus necessitating a larger degree of evaporation to reach saturation, but also to lower initial salinity and isotopically lighter water isotope composition. Thus, such addition of ion-free meteoric waters to the initial fluid changes neither the salinity ($115 g kg À1 ) or the water isotope composition (d 18 O $ 4‰ vs SMOW) of the gypsum saturation point (Fig. 4).

In these conditions, the mixing-evaporation model calculates a strong under-saturation with respect to gypsum (0.06 < X Gypsum < 0.32) at the salinity corresponding to that of the fluid inclusions of low-salinity Primary Lower Gypsum and Upper Gypsum deposits (red lines in Fig. 4b). We carried out additional model runs to test if the presence of dissolved Ca 2+ and SO 4 2À in continental runoff depressed the salinity at the point of gypsum saturation (dotted blue lines in Fig. 4). First, we repeated the above-mentioned model runs adding 5 mM dissolved Ca 2+ and SO 4 2À in the continental runoff end member. This set of model runs shows that the presence of dissolved Ca 2+ and/or SO 4 2À in the freshwater endmember depresses the salinity at gypsum saturation (Fig. 4b): as the runoff/ocean water ratio f R increases from 0.4 to 0.9, this effect increases in magnitude.

We considered two main sources of poorly-saline, Ca 2+and SO 4 2À -bearing fluids pertinent to our Mediterranean case ( rivers, based on recent studies suggesting that the Paratethys was a source of water for the Mediterranean Sea during the Messinian Salinity Crisis (Grothe et al., 2014(Grothe et al., , 2020;;Marzocchi et al., 2016;van Baak et al., 2017), and supposing that the Paratethys was a freshwater lake; (ii) modern Caspian Sea water, based on the idea (Grothe et al., 2020) that the Paratethys flowing into the Mediterranean during the Messinian Salinity Crisis had a chemical composition similar to that of the modern Caspian Sea (salinity $ 12 g kg À1 and dissolved Ca 2+ and SO 4 2À similar to modern seawater). In this set of model runs, to track the runoff/ocean water mixing ratio with geochemically measurable quantities, we consider the isotope tracers 87 Sr/ 86 Sr and d 34 S SO4 . The use of Sr isotopes to track mixing of Atlantic Ocean water with continental runoff is straightforward because Sr is not involved in biogeochemical processes and Sr isotopes do not fractionate during gypsum precipitation (Flecker et al., 2002;Topper et al., 2011;Grothe et al., 2020). The 87 Sr/ 86 Sr composition of the continental runoff end members is summarized in Table 2. The 87 Sr/ 86 Sr composition calculated by the model is compared to the 87 Sr/ 86 Sr of gypsum from the Primary Lower Gypsum and Upper Gypsum units presented in Roveri et al. (2014).

In addition to mixing between the oceanic and continental SO 4 2À sources, the d 34 S SO4 of dissolved SO 4

2À

, however, is affected by biogeochemical processes of the S cycle (sulfate reduction, sulfide oxidation, disproportionation...) taking place in sediments or water column during the Messinian Salinity Crisis (Pierre, 1982a,b;Lu, 2006;[START_REF] Garcia-Veigas | Geochemical indicators in Western Mediterranean Messinian evaporites: Implications for the salinity crisis[END_REF]. Although a large S-isotope fractionation (5 < e SR ‰ < 70; Sim et al., 2011;Leavitt et al., 2013)i s associated to SO 4 2À -consumption via sulfate reduction, resulting in a 34 S-enrichment of the residual SO 4 2À pool, successive biogeochemical production of SO 4 2À via re-oxidation of H 2 S can at most, in the case of a geochemically closed system, bring the d 34 S SO4 of the dissolved SO 4 2À pool back to its original, unreacted value (Aller et al., 2010;[START_REF] Mills | Geochemical evidence for cryptic sulfur cycling in salt marsh sediments[END_REF]Blonder et al., 2017). Therefore, the biogeochemical sulfur cycle rarely results in a net decrease of d 34 S of dissolved SO 4 2À and the d 34 S SO4 of Primary Lower Gypsum and Upper Gypsum deposits can be used to track the 34 S-depleted continental SO 4 2À source introduced into the Mediterranean basin by continental runoff or via the Paratethys. To this end we compiled available published d 34 S SO4 data from Primary Lower Gypsum and Upper Gypsum deposits and, together with the data presented in this work (Fig. 5), we defined a lower boundary for the d 34 S SO4 of Primary Lower Gypsum and Upper Gypsum that we use as a constraint to the mixing proportions between oceanic and continental SO 4

2À imposed in the model calculations.

Fig. 6 shows how the salinity at gypsum saturation, the 87 Sr/ 86 Sr and the d 34 S SO4 of Mediterranean Sea waters are affected by the runoff/ocean water mixing ratio, the concentration of riverine dissolved Ca 2+ ,S O 4 2À and Sr 2+ , and the isotope composition of riverine dissolved Sr ( 87 Sr/ 86 Sr). The isotope composition of riverine dissolved SO 4 2À (d 34 S SO4 ) was set equal to 8‰ vs CDT, corresponding to the average d 34 S SO4 of the Volga and Danube rivers [START_REF] Mekhtieva | Isotope Balance of Sulfur in the Caspian Sea[END_REF])( d 34 S SO4 for Mediterranean large rivers is not documented). The evaporation model complemented by these isotope systems was run 3000 times with different combinations of the forcing parameters listed above. Solutions to the model runs are presented in six diagrams corresponding to runoff/ocean water mixing ratios equal to 0.9 (Fig. 6a andb), 0.6 (Fig. 6c andd) and 0.4 (Fig. 6e andf). Isolines of salinity at gypsum saturation are represented as red lines in Fig. 6a,c and e and are read as a function of the riverine dissolved SO 4 2À (y axis) and Ca 2+ (x axis) concentration. In the same diagrams the d 34 S SO4 of dissolved sulfate in the mixture is represented by horizontal dashed -bearing runoff). In (a), model curves are compared to the fluid-inclusion salinity and d 18 O H2O of parent waters of gypsum deposits from Primary Lower Gypsum cycle 6 from the Sorbas basin (blue circles) [START_REF] Evans | Coupled measurements of delta O-18 and delta D of hydration water and salinity of fluid inclusions in gypsum from the Messinian Yesares Member, Sorbas Basin (SE Spain)[END_REF]. Notes: f R is the fraction of continental runoff mixed with Atlantic ocean waters; evaporative conditions for the salt works scenario are based on atmospheric measurements at the Vaccare `s Lagoon site (Camargue area, southern France; [START_REF] Delattre | Deuterium excess in the atmospheric water vapour of a Mediterranean coastal wetland: regional vs. local signatures[END_REF]; green arrows in panel (b) show the average salinity of fluid inclusions in the Piedmont basin (PB) [START_REF] Natalicchio | Did Late Miocene (Messinian) gypsum precipitate from evaporated marine brines? Insights from the Piedmont Basin (Italy)[END_REF], the Catanzaro Trough (CT) [START_REF] Costanzo | Messinian twinned selenite from the Catanzaro Trough, Calabria, Southern Italy: field, petrographic and fluid inclusion perspectives[END_REF], the Sorbas basin (SB) [START_REF] Evans | Coupled measurements of delta O-18 and delta D of hydration water and salinity of fluid inclusions in gypsum from the Messinian Yesares Member, Sorbas Basin (SE Spain)[END_REF] and the Vena del Gesso basin (VdG) (this work).

G. Aloisi et al. / Geochimica et Cosmochimica Acta xxx (2022) blue lines; the blue shaded area represents the range of salinities measured from fluid inclusions in gypsum from the Primary Lower Gypsum and Upper Gypsum units [START_REF] Natalicchio | Did Late Miocene (Messinian) gypsum precipitate from evaporated marine brines? Insights from the Piedmont Basin (Italy)[END_REF][START_REF] Evans | Coupled measurements of delta O-18 and delta D of hydration water and salinity of fluid inclusions in gypsum from the Messinian Yesares Member, Sorbas Basin (SE Spain)[END_REF][START_REF] Costanzo | Messinian twinned selenite from the Catanzaro Trough, Calabria, Southern Italy: field, petrographic and fluid inclusion perspectives[END_REF]; the red dashed box represents the dissolved Ca 2+ and SO 4 2À concentration in main rivers entering the Mediterranean, Black and Caspian seas (Table 2); the vertical orange and green bars represent the d 34 S of dissolved SO 4 2À deduced form isotope measurements of gypsum from the Primary Lower Gypsum and Upper Gypsum units (Fig. 5), while the red dot corresponds to the dissolved Ca 2+ and SO 4 2À in the modern Caspian Sea. In Fig. 6b, d andf, the blue contours represent the calculated 87 Sr/ 86 Sr of dissolved Sr 2+ as a function of the riverine dissolved Sr 2+ concentration (x-axis) and isotope composition ( 87 Sr/ 86 Sr) (y-axis); the orange and light blue shaded areas represent the 87 Sr/ 86 Sr range of Primary Lower Gypsum and Upper Gypsum deposits, respectively (Roveri et al., 2014); the red dots represent the Sr flux-weighted dissolved Sr concentration and 87 Sr/ 86 Sr of runoff considering only the Nile and the Rhone rivers (Nile/Rhone), only rivers draining into the modern Black ad Caspian seas (BS/CS rivers), and all of these sources (All rivers) (Table 2).

Dissolved Ca 2+ and SO 4 2À concentrations in the range of modern Mediterranean and Caspian/Black Sea rivers (red dashed box in Fig. 6a, c ande), coupled to a very high runoff fraction (f R = 0.9), result in a salinity at gypsum saturation that overlaps with the highest salinities (45 < S g kg À1 < 78.2) deduced for Primary Lower Gypsum and Upper Gypsum deposits based on gypsum fluid inclusions [START_REF] Costanzo | Messinian twinned selenite from the Catanzaro Trough, Calabria, Southern Italy: field, petrographic and fluid inclusion perspectives[END_REF][START_REF] Evans | Coupled measurements of delta O-18 and delta D of hydration water and salinity of fluid inclusions in gypsum from the Messinian Yesares Member, Sorbas Basin (SE Spain)[END_REF][START_REF] Natalicchio | Did Late Miocene (Messinian) gypsum precipitate from evaporated marine brines? Insights from the Piedmont Basin (Italy)[END_REF] this work) (blue area in Fig. 6a, c ande). With these Ca 2+ and SO 4 2À concentrations, the salinity at saturation remains higher than the average salinity of Primary Lower Gypsum and Upper Gypsum fluid inclusions from the Piedmont basin (16 g kg À1 ) ( Natalicchio et al., 2014), the Catanzaro Trough (27 g kg À1 ) [START_REF] Costanzo | Messinian twinned selenite from the Catanzaro Trough, Calabria, Southern Italy: field, petrographic and fluid inclusion perspectives[END_REF] and the Sorbas Basin (32 g kg À1 ) [START_REF] Evans | Coupled measurements of delta O-18 and delta D of hydration water and salinity of fluid inclusions in gypsum from the Messinian Yesares Member, Sorbas Basin (SE Spain)[END_REF]. Only the average salinity of the Vena del Gesso fluid inclusions (52.2 g kg À1 ; this work) is compatible with f R = 0.9 and the dissolved Ca 2+ and SO 4 2À concentration of modern Mediterranean and Caspian/Black Sea rivers. With these riverine fresh water sources, however, the final mixture has 87 Sr/ 86 Sr comparable to that of Upper Gypsum deposits (Fig. 6b), which are more 87 Sr-depleted than the Vena del Gesso Primary Lower Gypsum deposits. On the other hand, imposing f R values lower than 0.9 (Fig. 6c ande), dissolved Ca 2+ and SO 4 2À concentrations in the range of modern Mediterranean and Caspian/Black Sea rivers result in salinities higher than those of Primary Lower Gypsum and Upper Gypsum fluid inclusions. On this basis, we can exclude that the small amounts of riverine dissolved Ca 2+ and SO 4 2À derived from weathering of carbonates and pyrite can be the trigger for the precipitation of low-salinity gypsum in Mediterranean marginal basins.

With a runoff component having the dissolved Ca 2+ and SO 4 2À composition of the modern Caspian Sea (red dot in Fig. 6a, c ande), the salinity at gypsum saturation is considerably depressed at all three f R values (14 < S g kg À1 < 66) and covers most of the range of salinities found in Primary Lower Gypsum and Upper Gypsum fluid inclusions [START_REF] Natalicchio | Did Late Miocene (Messinian) gypsum precipitate from evaporated marine brines? Insights from the Piedmont Basin (Italy)[END_REF][START_REF] Evans | Coupled measurements of delta O-18 and delta D of hydration water and salinity of fluid inclusions in gypsum from the Messinian Yesares Member, Sorbas Basin (SE Spain)[END_REF][START_REF] Costanzo | Messinian twinned selenite from the Catanzaro Trough, Calabria, Southern Italy: field, petrographic and fluid inclusion perspectives[END_REF]. However, since the modern Caspian Sea has an elevated dissolved SO 4 2À concentration (30 mM; Table 2), the resulting d 34 S SO4 of the seawater-runoff mixture is low and varies from 9‰ vs CDT (f R = 0.9) to 16‰ vs CDT (f R = 0.4). Therefore, for all considered values of f R , the d 34 S SO4 of the resulting seawater-runoff mixture is considerably lower than that deduced from d 34 S SO4 measurements of both Primary Lower Gypsum and Upper Gypsum deposits (20.15 < d 34 S SO4 vs CDT < 23.7) (Fig. 6a, c ande). As discussed in Section 6.4, these results provide a quantitative basis for evaluating the influence of the Paratethys in the formation of low-salinity gypsum (Grothe et al., 2020).

INTERPRETATION AND DISCUSSION

Primary versus diagenetic origin of the ''low-salinity geochemical signal

Based on the absence of macro-and microscale evidence for recrystallization of gypsum (e.g. Fig. 7), we can exclude that pervasive dissolution-reprecipitation processes or dehydration/rehydration cycles have reset the gypsum crystallization water isotope composition (see supplementary material for a full discussion of gypsum diagenesis). Indeed, a simple isotope mass balance calculation (Table 3) implies that if the measured isotope composition of gypsum samples was the result of diagenetic alteration of an original, evaporitic gypsum (with crystallization water d 18 O $ 6.5‰ vs SMOW) under the influence of meteoric fluids (À9<d 18 O‰ vs SMOW < À5.11; Table 3), then from 5% to 87% of the gypsum in the analyzed samples would have suffered intense post-depositional alteration. Petrographic analyses of the crystals carried out on 15 polished slabs and 30 thin slices (see supplementary methods), however, do not show any evidence for a pervasive recrystallization process, but rather suggest the perfect preservation of delicate pristine features of the crystals such as filamentous microfossils and diatom remains (e.g., Pellegrino et al., 2021).

As for the fluid inclusion salinity signal, it can potentially be reset by leaking of the fluid inclusions and partial refilling with meteoric waters, without affecting the isotope composition of gypsum crystallization water. However, we are confident that the studied samples are mostly not affected by such a contamination for the following reasons:

(1) we have measured fluid inclusions of different sizes (10-60 lm wide, with an area approximately ranging from 100 to 3500 lm 2 ), without optical evidence of modification (i.e. irregular shape of the inclusions, fractures); (2) we systematically excluded fluid inclusions located near fluid inclusion families of clear secondary origin (see supplementary methods for petrographic criteria used to discriminate primary from secondary fluid inclusions); [START_REF] Lein | Isotopic composition of sulfur compounds in sediments of the Pacific[END_REF] we have carried out an exploratory cathodoluminescence survey of what we interpret as primary and secondary fluid inclusions based on petrographic evidence. We observed no difference of luminescence between the primary fluid inclusions and the surrounding gypsum. In contrast, different generations of secondary fluid inclusions (both trails along fractures and fluid inclusion families placed along the main gypsum cleavage plains) show occasionally a pale to bright orange Fig. 5. d 34 S SO4 composition of the dissolved sulfate ion in equilibrium with Primary Lower Gypsum and Upper Gypsum deposits formed in the marginal basins of the Mediterranean during the Messinian Salinity Crisis. Red crosses are the average d 34 S SO4 values for given marginal basins (gray boxes) and for all the complete Primary Lower Gypsum and Upper Gypsum datasets (orange box, n = 281 for Primary Lower Gypsum and n = 94 for Upper Gypsum). The lower and upper bounds of the boxes correspond to the 1st and 3rd quartile; the horizontal line inside the boxes represents the median; vertical lines (whiskers) extend to minimum and maximum values, excluding outliers (represented as small circles). The d 34 S SO4 value from the Messinian ocean (22‰ vs CDT) is taken from Masterson et al. (2016). Data sources are: Piedmont basin (this work, Supplementary Tables 678); Caltanissetta basin (this work, Supplementary Tables 234; [START_REF] Garcia-Veigas | Geochemical indicators in Western Mediterranean Messinian evaporites: Implications for the salinity crisis[END_REF]; Sorbas basin [START_REF] Playa | Marine to non-marine sedimentation in the upper Miocene evaporites of the Eastern Betics, SE Spain: sedimentological and geochemical evidence[END_REF][START_REF] Evans | Coupled measurements of delta O-18 and delta D of hydration water and salinity of fluid inclusions in gypsum from the Messinian Yesares Member, Sorbas Basin (SE Spain)[END_REF][START_REF] Garcia-Veigas | Geochemical indicators in Western Mediterranean Messinian evaporites: Implications for the salinity crisis[END_REF]; Nijar basin (Lu et al., 2001); Vena del Gesso (Supplementary Table 5 and (Lugli et al., 2007); Mallorca [START_REF] Garcia-Veigas | Geochemical indicators in Western Mediterranean Messinian evaporites: Implications for the salinity crisis[END_REF]; Feos (Lu et al., 2001); Polemi basin (Pierre, 1982a,b) S3) suggesting the presence of carbonate phases precipitated from postdepositional groundwater fluids. Overall, these observations confirm that the studied primary fluid inclusions still retain their primary chemical composition. These results have to be weighed in the light of a recent work by Bigi et al. (2022) that applies a new microthermometric protocol for fluid inclusion analysis of Messinian gypsum samples, arguing against the low-salinity scenario for marginal basin Messinian gypsum deposits. Bigi et al. (2022) propose that the more reliable salinity data, falling in the classical gypsum precipitation salinity field (salinity > 110 g kg À1 ), are from the smallest fluid inclusions (with area < 700 lm 2 ), while larger fluid inclusions, that include those bearing the low-salinity signal, are more likely affected by leaking and refilling processes. However, both in [START_REF] Natalicchio | Did Late Miocene (Messinian) gypsum precipitate from evaporated marine brines? Insights from the Piedmont Basin (Italy)[END_REF], as well as in the present study, we have measured numerous primary fluid inclusions falling in the <700 lm 2 size class that have salinities well below a typical evaporated, gypum-saturated seawater. Further, the new microthemometric protocol proposed by Bigi et al (2022) implies a strong thermal stretching consisting in several cooling-heating cycles (thermal shock, from À100 to +120 °C) with the aim to nucleate a small vapor bubble within the inclusions that reduce their metastabele behavior during microthemometric measurements. Such an approach largely differs from the weak stretching methodology used in this work which consisted in a single cooling-heating cycle (À90 to +30 °C), as suggested by previous studies [START_REF] Attia | Middle Miocene gypsum, Gulf of Suez: Marine or nonmarine?[END_REF][START_REF] Natalicchio | Did Late Miocene (Messinian) gypsum precipitate from evaporated marine brines? Insights from the Piedmont Basin (Italy)[END_REF], in order to promote the nucleation of small bubbles while minimizing the loss of fluid from the inclusion. Indeed, especially after stretching, the fluid can be loss by diffusion from the inclusion, even in cases of low strain rate (e.g. [START_REF] Bodnar | Revised Equation and Table for Determining the Freezing-Point Depression of H 2 O-Nacl Solutions[END_REF]. We suspect that a strong thermal shock applied to soft gypsum crystals may cause an excessive stretching (i.e. increase in volume) with an irreversible re-equilibration of fluid inclusions. Experimental studies from aqueous fluid inclusions in quartz [START_REF] Hall | Preferential Water-Loss from Synthetic Fluid Inclusions[END_REF] showed that the water loss by diffusion and mechanical micro-stretching resulted in an increase of the fluid inclusion salinity, despite the elevated hardness of this mineral. We therefore conclude that the fluid inclusion salinity and the isotopic composition of gypsum hydration water of Primary Lower Gypsum and Upper Gypsum deposits presented in this work represent a primary geochemical signal.

Updating existing evidence for low-salinity gypsum deposits during the Messinian Salinity Crisis

The new fluid inclusion data from the Vena del Gesso expands the existing evidence for gypsum from the Primary Lower Gypsum and Upper Gypsum units precipitated from low-salinity water masses [START_REF] Natalicchio | Did Late Miocene (Messinian) gypsum precipitate from evaporated marine brines? Insights from the Piedmont Basin (Italy)[END_REF][START_REF] Evans | Coupled measurements of delta O-18 and delta D of hydration water and salinity of fluid inclusions in gypsum from the Messinian Yesares Member, Sorbas Basin (SE Spain)[END_REF][START_REF] Costanzo | Messinian twinned selenite from the Catanzaro Trough, Calabria, Southern Italy: field, petrographic and fluid inclusion perspectives[END_REF]. Of the samples for which we do not have fluid inclusion data (Eraclea Minoa, Santa Elisabetta, Gibliscemi, Arnulfi and Pollenzo sections), some indications of formation from low-salinity parent waters come from the hydration water isotope composition. An intuitive -but wrong -interpretation, is that Messinian gypsum parent water isotope compositions significantly lighter than those of modern salt work gypsum [START_REF] Fontes | Fractionnement isotopique dans l'eau de cristallisation du sulfate de calcium[END_REF][START_REF] Fontes | Fractionnement isotopique de l'hydroge `ne dans l'eau de cristallisation du gypse[END_REF]Longinelli, 1979;[START_REF] Evans | Coupled measurements of delta O-18 and delta D of hydration water and salinity of fluid inclusions in gypsum from the Messinian Yesares Member, Sorbas Basin (SE Spain)[END_REF] imply that Messinian gypsum was formed from low-salinity (<115 g kg À1 ) water masses. Indeed, all gypsum parent waters from the considered marginal basins are isotopically lighter than those of gypsum formed by evaporation of seawater in modern Mediterranean Sea salt works (Fig. 2).

Our evaporation calculations, however, show that there is a strong continental influence on the hydration water isotope composition of gypsum formed in modern Mediterranean salt works (Fig. 2f). Both the Vaccare `s Lagoon and the Aigues Mortes salt works are located 4-10 km inland from the coast, and experience more continental atmospheric conditions than those of the open Mediterranean Sea. Typically, atmospheric relative humidity and wind speed are lower over land than over the open ocean [START_REF] Archer | Evaluation of global wind power[END_REF]. Also, the water temperature imposed in the model to represent that of the Aigues Mortes salt work summer waters (29 °C) is several degrees higher than the Mediterranean Sea summer surface temperature (22-27 °C;[START_REF] Pastor | Sea Surface Temperature in the Mediterranean: Trends and Spatial Patterns (1982-2016)[END_REF].

Consistently, when average yearly or July evaporation conditions over the open Mediterranean Sea -that are more humid and windier that those at the Vaccare `s Lagoon -are 3 Fig. 6. Modelled salinity and isotope composition (d 34 S SO4 and 87 Sr/ 86 Sr) of gypsum deposits precipitating from a mixture of seawater and river water with variable dissolved Ca 2+ and SO 4 2À concentration (panels a,c and e) and Sr 2+ concentrations and 87 Sr/ 86 Sr ratios (panels b, d and f). The graphs are drawn for a fraction of river water in the mixture equal to 0.9 (panels a and b), 0.6 (panels c and d) and 0.4 (panels e and f). In panels a, c and e, red contours represent the calculated salinity at gypsum saturation while the blue shaded area represents the range of salinities measured from fluid inclusions in gypsum from the Primary Lower Gypsum and Upper Gypsum units [START_REF] Natalicchio | Did Late Miocene (Messinian) gypsum precipitate from evaporated marine brines? Insights from the Piedmont Basin (Italy)[END_REF][START_REF] Evans | Coupled measurements of delta O-18 and delta D of hydration water and salinity of fluid inclusions in gypsum from the Messinian Yesares Member, Sorbas Basin (SE Spain)[END_REF][START_REF] Costanzo | Messinian twinned selenite from the Catanzaro Trough, Calabria, Southern Italy: field, petrographic and fluid inclusion perspectives[END_REF]; horizontal dotted lines represent the calculated d 34 S of dissolved SO 4 2À while the vertical orange and green bars represent the d 34 S of dissolved SO 4 2À deduced form isotope measurements of Primary Lower Gypsum and Upper Gypsum deposits (see text for references); the red dotted area represents the range of modern dissolved Ca 2+ and SO 4 2À in the main rivers entering the Mediterranean, Black and Caspian seas (Table 2), while the red dot corresponds to the dissolved Ca 2+ and SO 4 2À in the modern Caspian Sea. In panels b, d and f, the blue contours represent the calculated 87 Sr/ 86 Sr of dissolved Sr; the orange and light blue shaded areas represent the 87 Sr/ 86 Sr range of gypsum from the Primary Lower Gypsum and Upper Gypsum units, respectively (see text for references); the red dots represent the Sr flux-weighted dissolved Sr concentration and 87 Sr/ 86 Sr ratio of runoff considering only the Nile and the Rhone rivers (Nile/ Rhone), only rivers draining into the modern Black ad Caspian seas (BS/CS rivers), and all of these sources (All rivers).

G. Aloisi et al. / Geochimica et Cosmochimica Acta xxx (2022) considered, the calculated isotope composition of water at the gypsum saturation point is up to several per mill lighter in 18 O and D than that of modern salt works (Fig. 2f). Thus, the parent water isotope composition of most of the samples from the Santa Elisabetta and Eraclea Minoa sections fall in the same range of d 18 O H2O values calculated at gypsum saturation by the model forced by average yearly Mediterranean Sea conditions. For this reason, the gypsum samples from these two sections are compatible with formation from the evaporation of seawater to a salinity of 115 g kg À1 and cannot, based on our data, be interpreted as low-salinity gypsum. On the other hand, gypsum samples from the Pollenzo and Arnulfi sections of the Piedmont Basin, and the most 18 O-depleted samples from the Gibliscemi and the Santa Elisabetta sections in the Caltanissetta basin, have parent waters more 18 O-depleted than those calculated by the model at the gypsum saturation point. These gypsum samples have been formed from water masses: (i) that contained an important admixture of 18 O-depleted continental runoff, and (ii) that underwent only a limited extent of evaporation. These conditions imply that these are low-salinity gypsum deposits. Altogether, four Mediterranean marginal basins -the Piedmont basin [START_REF] Natalicchio | Did Late Miocene (Messinian) gypsum precipitate from evaporated marine brines? Insights from the Piedmont Basin (Italy)[END_REF], the Sorbas basin [START_REF] Evans | Coupled measurements of delta O-18 and delta D of hydration water and salinity of fluid inclusions in gypsum from the Messinian Yesares Member, Sorbas Basin (SE Spain)[END_REF], the Catanzaro Trough [START_REF] Costanzo | Messinian twinned selenite from the Catanzaro Trough, Calabria, Southern Italy: field, petrographic and fluid inclusion perspectives[END_REF] and the Vena del Gesso basin (this work) -contain gypsum deposits that, based on fluid inclusion salinity estimates, were formed from water masses that were far less saline than the salinity at which gypsum starts to precipitate from evaporating seawater ($110 g kg À1 )( Table 2). In addition, gypsum from the Pollenzo and Arnulfi section of the Piedmont basin, and the most 18 O-depleted samples from the Gibliscemi and Santa Elisabetta sections, can be interpreted as low-salinity gypsum deposits based on the isotope composition of gypsum hydration water. The overwhelming majority of evidence for the occurrence of lowsalinity gypsum comes from Primary Lower Gypsum, rather than Upper Gypsum deposits.

Recycling of older gypsum deposits

The apparent contradiction of gypsum deposits bearing geochemical signatures that indicate precipitation from a poorly saline water mass has been explained by the leaching of older, exposed gypsum deposits and the precipitation of the released SO 4 2À and Ca 2+ ions in adjacent basins [START_REF] Attia | Middle Miocene gypsum, Gulf of Suez: Marine or nonmarine?[END_REF][START_REF] Ayora | Modeling the sulfur and oxygen isotopic composition of sulfates through a halite-potash sequence: Implications for the hydrological evolution of the Upper Eocene Southpyrenean Basin[END_REF][START_REF] Cendo ´n | The importance of recycling processes in the Middle Miocene Badenian evaporite basin (Carpathian foredeep): palaeoenvironmental implications[END_REF]. The recycling process can be easily tracked when the preexisting gypsum is much older and bears a d 34 S SO4 signature that is distinct (usually 34 S-depleted; see Fig. 3) compared to the neo-formed gypsum deposit. This is the case of gypsum from the Upper Eocene South Pyrenean Basin [START_REF] Ayora | Modeling the sulfur and oxygen isotopic composition of sulfates through a halite-potash sequence: Implications for the hydrological evolution of the Upper Eocene Southpyrenean Basin[END_REF], formed by recycling of Upper Triassic 34 S-depleted gypsum deposits. This recycling process led to the deposition of gypsum with a very low d 34 S SO4 (12.95 < d 34 S SO4 ‰ vs CDT < 19.27) compared to primary gypsum formed in the same basin by precipitation from seawater (average d 34 S SO4 = 21.6). Recycling of Upper Triassic evaporites was also reported for the deposition of the upper Miocene gypsum of intramontane Neogene Betic basins, such as the Lorca and Fortuna basins [START_REF] Playa | Marine to non-marine sedimentation in the upper Miocene evaporites of the Eastern Betics, SE Spain: sedimentological and geochemical evidence[END_REF][START_REF] Garcia-Veigas | Late Miocene evaporite geochemistry of Lorca and Fortuna basins (Eastern Betics, SE Spain): Evidence of restriction and continentalization[END_REF]. For the Primary Lower Gypsum and Upper Gypsum deposits studied here, recycling of older gypsum can be, however, ruled out, based on the late Miocene d 34 S SO4 signature of these deposits.

The effect of recycling on the d 34 S SO4 of gypsum is undetectable when the recycling takes place intra-basinally and thus the original, weathered primary gypsum has the same age, and the same d 34 S SO4 , as the neo-precipitated gypsum. This is the case described by [START_REF] Attia | Middle Miocene gypsum, Gulf of Suez: Marine or nonmarine?[END_REF] to explain primary gypsum deposits from the Middle Miocene of Egypt (Gulf of Suez) with d 34 S SO4 (21.9-23.6‰ vs CDT) indicating a marine SO 4 2À source but low-salinity fluid inclusions. A similar scenario was proposed for the low salinity Badenian [START_REF] Cendo ´n | The importance of recycling processes in the Middle Miocene Badenian evaporite basin (Carpathian foredeep): palaeoenvironmental implications[END_REF] and Messinian primary marine gypsum [START_REF] Evans | Coupled measurements of delta O-18 and delta D of hydration water and salinity of fluid inclusions in gypsum from the Messinian Yesares Member, Sorbas Basin (SE Spain)[END_REF], including the Piedmont gypsum studied here (Banengo section, [START_REF] Natalicchio | Did Late Miocene (Messinian) gypsum precipitate from evaporated marine brines? Insights from the Piedmont Basin (Italy)[END_REF]. However, no geological evidence of coeval marine evaporitic gypsum that was leached to provide the ions necessary for the precipitation of the observed low salinity gypsum has been reported, making this explanation unlikely and not supported by field data. Moreover, the majority of the oldest evaporitic deposits that we analyzed, i.e those from the 1st Primary Lower Gypsum cycle from the Monticino (Vena del Gesso) and Banengo and Arnulfi (Piedmont Basin) sections have fluid inclusion salinity or hydration water isotopes that indicate formation from low-salinity water masses (most of the gypsum hydration water isotopes from the S. Elisabetta section are inconclusive as to the salinity of the parent waters, see above): the hypothetical, original marine gypsum deposits that would have been leached are missing.

Ca-and SO 4 -rich Paratethys inflow during the Messinian Salinity Crisis

Based on 87 Sr/ 86 Sr isotope evidence for a connection between the Paratethys and the Mediterranean Sea starting from 6.1 Ma, Grothe et al. (2020) propose that brackish water flowing into the Mediterranean Sea from the (Grothe et al., 2020), has a salinity ranging from 10 to 15 g kg À1 , but dissolved Ca 2+ and SO 4 2À concentrations similar to modern seawater ($8 mol kg À1 and $30 mol kg À1 , respectively). By evaporating this water, and a modified version without dissolved Na + and Cl À , gypsum saturation is attained at 45 g kg À1 and 14 g kg À1 , respectively (Grothe et al., 2020).

Several lines of evidence, however, suggest that the existence of such a Ca 2+ -and SO 4 2À -rich Paratethyan source is unlikely. First, if Paratethys waters flowing into the Mediterranean Sea and into its marginal basins did not mix with the underlying Mediterranean water mass, as the evaporation calculations of Grothe et al. (2020) suppose, then they should have retained the 87 Sr/ 86 Sr signature of the Paratethys ($0.7084-0.7085; Grothe et al. 2020). This is not the case because gypsum from the Primary Lower Gypsum unit shows a uniform 87 Sr/ 86 Sr signature that is more radiogenic (0.708640 < 87 Sr/ 86 Sr < 0.709024) than the Paratethys Sr source, implying mixing with Atlantic-derived Sr ( 87 Sr/ 86 Sr $ 0.7090; [START_REF] Mcarthur | Strontium Isotope Stratigraphy: LOWESS Version 3: Best fit to the Marine Sr-Isotope curve for 0-509 Ma and accomanying lookup table for deriving numerical age[END_REF]. In addition to changing the 87 Sr/ 86 Sr, mixing of the Paratethys source with Atlantic Ocean waters would have increased the salinity of the mixture. Second, as our model calculations show (Section 5.2), lack of mixing between the Paratethys input and Atlantic waters would result in a strongly 34 S-depleted dissolved sulfate ion similar to the one of the rivers flowing into the Black and Caspian Seas (0.7 < d 34 S SO4 ‰ vs CDT < 7.6; Table 3), or the modern Caspian Sea (d 34 S SO4 = 8.7‰ vs CDT). However, this is not the case because all Primary Lower Gypsum and Upper Gypsum deposits have d 34 S SO4 values indicating a marine SO 4 2À source (20.4 < d 34 S SO4 ‰ vs CDT < 23.7; Fig. 3).

In our modeling exercise discussed in section 5.2, the least well-constrained forcing parameter is the dissolved Ca 2+ concentration of the non-oceanic source. The modern rivers flowing into the Mediterranean, Black and Caspian Seas all have dissolved Ca 2+ concentrations 2m M (Table 3). With these dissolved Ca 2+ concentrations, and the constraints imposed by the 87 Sr/ 86 Sr and d 34 S SO4 signature of gypsum deposits, the model always calculates a salinity at gypsum saturation greater than 60 g kg À1 (Fig. 6). Since it is likely that dissolved Ca 2+ concentrations in Messinian rivers feeding the Mediterranean, Black and Caspian Seas during the Messinian Salinity Crisis were not radically different from the modern values, we can exclude that Ca 2+ -and SO 4 2À -bearing continental runoff is responsible for the formation of low-salinity gypsum. The only alternative, potentially Ca-rich, brackish water source is the Paratethys. Accepting Sr-isotope evidence suggesting a hydrological connection between the Mediterranean Sea and the Paratethyan Black Sea and Caspian domains during the Messinian Salinity Crisis (Grothe et al., 2020), the idea that the modern chemical composition of the Caspian Sea -brackish, but Ca 2+ -and SO 4 2À -rich -is representative of the Paratethys during the Messinian Salinity Crisis needs to be evaluated.

The Caspian Sea has been an endorheic basin during the past $3 Myr [START_REF] Palcu | The end of the Great Khersonian Drying of Eurasia: Magnetostratigraphic dating of the Maeotian transgression in the Eastern Paratethys[END_REF]. Its modern chemical composition has thus been acquired during 3 Myr of ion addition by rivers and ion extraction by chemical and biogeochemical processes [START_REF] Berner | Global Environment: Water, Air and Geochemical Cycles[END_REF]: its high Ca 2+ and SO 4 2À concentrations are the consequence of isolation. Prior to the connection of the Paratethys to the Mediterranean (before 6.1 Ma) a few 87 Sr/ 86 Sr data points from the Black and Caspian Sea domains indicate that 87 Sr/ 86 Sr in the Paratethys was low ($0.708025) relative to coeval oceanic and Mediterranean Sea 87 Sr/ 86 Sr ($0.709). This low Paratethyan 87 Sr/ 86 Sr signature is very similar to the Sr flux-weighted 87 Sr/ 86 Sr of modern Sr riverine input into the Black and Caspian seas ([Sr 2+ ] = 4.92 * 10 À6 mol/litre; 87 Sr/ 86 Sr = 0.708215). This supports the idea that also prior to 6.1 Ma the Paratethys was a large endoreic water mass that potentially had high dissolved Ca 2+ and SO 4 2À concentration similar to that of the modern Caspian Sea.

At roughly 6.1 Ma, 87 Sr/ 86 Sr in the Paratethyan domain rises sharply reaching values similar to coeval Mediterranean water, suggesting a connection with the Mediterranean (confirmed by eastern Mediterranean biomarker studies; Vasiliev et al. (2019), and then decrease to about 0.7084-0.7085 during the whole of the Messinian Salinity Crisis. These intermediate 87 Sr/ 86 Sr Messinian Salinity Crisis values are taken to represent mixing between the Paratethyan and Mediterranean Sea Sr (Grothe et al., 2020). Supposing that prior to the 6.1 Ma connection the chemical composition of the Paratethys was similar to that of the modern Caspian Sea [START_REF] Clauer | Fluctuations of Caspian Sea level: Beyond climatic variations?[END_REF]Grothe et al., 2020), could this ''endorheic" composition of the Paratethys have persisted throughout the 130 kyr preceding the Messinian Salinity Crisis onset when the Paratethys and the Mediterranean Sea were connected? The order-of-magnitude calculation that follows excludes this possibility.

The volume of the Paratethys at the end of the Miocene is estimated to have been between 0.5 and 1 Â 10 6 km 3 (de la Vara et al., 2016). With a late-Miocene Paratethys freshwater excess equal 0.02 Sv (=0.02 Â 10 6 m 3 /s) (Marzocchi et al., 2016), the whole water volume of the Paratethys would have been replaced 80 to 160 times in 130 kyr that preceded the Messinian Salinity Crisis onset. The endorheic character of Caspian Sea chemistry, acquired prior to the connection with the Mediterranean, would therefore have been washed away before the onset of the Messinian Salinity Crisis. Should the Paratethys-Mediterranean connection have been a one-way flow of water directed to the Mediterranean, then the Paratethys would have been transformed into a lake having the chemical composition corresponding to the average composition of the riverine input. If, instead, as Sr-isotopes (Grothe et al., 2020) Ca 2+ concentration higher than that of the rivers that fed the Paratethys (<2 mM); they also explain why gypsum from the Primary Lower Gypsum and Upper Gypsum units lacks a continental, strongly 34 S-depleted isotope signature.

6.5. Does the biogeochemical sulfur cycle play a role in the formation of low-salinty gypsum?

In this section we will develop on the idea originally proposed by Dela Pierre et al. (2014Pierre et al. ( , 2015)), and successively by Grothe et al. (2020), that the biogeochemical sulfur cycle is involved in the formation of Mediterranean low-salinity gypsum deposits.

Biogeochemically active Messinian Salinity Crisis marginal basins

Evidence from a large number of studies using complementary approaches, especially from the Primary Lower Gypsum unit, suggests that low-salinity gypsum deposits accumulated in biogeochemically active marginal basins. Both bottom grown selenite crystals and the marly interbeds of the Primary Lower Gypsum unit deposits contain fossils of densely-packed filamentous organisms (''spaghetti-like" structures) (Vai and Ricci Lucchi, 2006;[START_REF] Panieri | Ribosomal RNA gene fragments from G. Aloisi et al[END_REF](Fig. 7), that are also found in carbonate and shale layers considered as the lateral, deep-basin time-equivalents of gypsum layers of the Primary Lower Gypsum unit [START_REF] Oliveri | Carbonate stromatolites from a Messinian hypersaline setting in the Caltanissetta Basin, Sicily: petrographic evidence of microbial activity and related stable isotope and rare earth element signatures[END_REF]Dela Pierre et al., 2012;Caruso et al., 2015;Natalicchio et al., 2017Natalicchio et al., , 2019)). Recently the filaments were interpreted as being the fossil remains of giant colorless sulfide-oxidizing bacteria (CSB) like Beggiatoa and Thioploca on the basis of morphological features (size, hollowness, curved shape and internal segmentation) (Schopf et al., 2012;Dela Pierre et al., 2014[START_REF] Mas | El registre estratigr`afic del Messini`a terminal i del Pliocè a l'illa de Mallorca[END_REF] and the presence of iron sulfide and polysulfide inclusions within the filaments. Such inclusions were interpreted to represent the product of early diagenetic transformations of sulfur globules stored within the cells (Dela Pierre et al., 2015). The sulfur globules correspond to an intermediate product of the oxidation of sulfide to sulfate and are a clade diagnostic feature of colorless sulfide-oxidizing bacteria [START_REF] Huettel | Vertical Migration in the Sediment-Dwelling Sulfur Bacteria Thioploca spp. in Overcoming Diffusion Limitations[END_REF][START_REF] Schulz | Big bacteria[END_REF][START_REF] Teske | The Genera Beggiatoa and Thioploca[END_REF]. The assignment of the filamentous fossils to this group of bacteria suggests that the basin floor where most of the Messinian gypsum grew was covered by a benthic assemblage of chemotrophic bacteria involved in an intense sulfur cycle. Besides, filamentous microfossilbearing gypsum crystals contain abundant nano-sized planktic diatoms that reflects conditions of high primary productivity in the upper water column. These diatoms represented a probable source of organic matter exploitable by benthic bacteria (Pellegrino et al., 2021). In addition, both gypsum and the interbedded marls contain organic-rich aggregates interpreted as marine snow floccules formed by aggregation of clay and organic material in the upper water column during episodes of eutrophication and phytoplankton blooms and deposited on an anoxic basin floor (Dela Pierre et al., 2014[START_REF] Mas | El registre estratigr`afic del Messini`a terminal i del Pliocè a l'illa de Mallorca[END_REF].

In the bottom grown selenite crystals of the Primary Lower Gypsum unit, filaments and floccules are entrapped in mm-thick turbid laminae that, in the re-entrant angle of the twinned crystals, alternate with limpid lamina typified by scarce or absent biogenic material (Fig. 7). Turbid laminae contain pyrite and dolomite grains (Dela Pierre et al., 2015) that are common by-products of microbial sulfate reduction. The lamination in the gypsum crystals is interpreted as the product of seasonal climate variability [START_REF] Orti | Selenite facies in marine evaporites: A review[END_REF]Dela Pierre et al., 2015;Reghizzi et al., 2018). During the humid season, increased continental runoff favored water column stratification, high primary productivity and bottom oxygen depletion, favoring organic matter preservation in the sediments. Organic matter degradation by bacterial sulfate reduction provided the hydrogen sulfide flux required for the growth of CSB (Dela Pierre et al., 2015). The arid season was instead dominated by water column mixing and oxygenation of the water column (Dela Pierre et al., 2015). As discussed in the next section, such a seasonal alternation between oxidized and reduced conditions might have selected the benthic ecosystem in favor of metabolically flexible CSB communities.

Supplementary evidence for biogeochemical activity comes from the study of molecular fossils extracted from gypsum layers and marl interbeds of the Primary Lower Gypsum unit. The molecular fossil inventory of gypsum beds (hydrocarbons, alcohols, carboxylic acids) confirms that the mazes of filamentous microfossil are benthic microbial assemblages dominated by CSB associated to sulfate reducing bacteria (Natalicchio et al., 2022). The overall bacterial and archeal lipid inventory suggest the existence of a stratified marine basin (Natalicchio et al., 2022). With regards to the marl interbeds, molecular fossils include: (i) lipds produced by organisms that live at the interface between oxic and anoxic waters in stratified basins -including anoxygenic phototrophic bacteria - [START_REF] Liaaen-Jensen | Chemistry of carotenoid pigments[END_REF][START_REF] Kleemann | Tetrahymanol from the Phototrophic Bacterium Rhodopseudomonas-Palustris -1st Report of a Gammacerane Triterpene from a Prokaryote[END_REF][START_REF] Damste | Molecular indicators for palaeoenvironmental change in a Messinian evaporitic sequence (Vena del Gesso, Italy). I: Variations in extractable organic matter of ten cyclically deposited marl beds[END_REF][START_REF] Kenig | Molecular Indicators for Paleoenvironmental Change in a Messinian Evaporitic Sequence (vena Del Gesso, Italy). 2. High-Resolution Variations in Abundances and C-13 Contents of Free and Sulfur-bound Carbon Skeletons in a Single Marl Bed[END_REF][START_REF] Rashby | Biosynthesis of 2-methylbacteriohopanepolyols by an anoxygenic phototroph[END_REF][START_REF] Wakeham | Microbial ecology of the stratified water column of the Black Sea as revealed by a comprehensive biomarker study[END_REF][START_REF] Eickhoff | Oxidation of Fe(II) leads to increased C-2 methylation of pentacyclic triterpenoids in the anoxygenic phototrophic bacterium Rhodopseudomonas palustris strain TIE-1[END_REF], implying that anaerobia impinged on the photic zone during the deposition of marl interbeds [START_REF] Tenhaven | Tetrahymanol, the most likely precursor of gammacerane, occurs ubiquitously in marine sediments[END_REF][START_REF] Schoell | Carbon isotopic composition of individual biomarkers in gilsonites (Utah)[END_REF][START_REF] Damste | Molecular indicators for palaeoenvironmental change in a Messinian evaporitic sequence (Vena del Gesso, Italy). I: Variations in extractable organic matter of ten cyclically deposited marl beds[END_REF], and (ii) lipids with C and N isotope compositions implying that N 2fixing diazotrophic cyanobacteria drove primary production during deposition of the marl interbeds (Isaji et al., 2019).

In search of a modern biogeochemical analogue

In modern marine environments, giant CSB like Beggiatoa and Thioploca thrive in organic-rich sediments such as in the Pacific Peru and the Atlantic Namibian continental margins [START_REF] Jorgensen | Distribution of colorless sulfur bacteria (Beggiatoa spp.) in a coastal marine sediment[END_REF][START_REF] Gallardo | Large benthic microbial communities in sulphide biota under Peru-Chile Subsurface Countercurrent[END_REF][START_REF] Mussmann | Phylogeny and distribution of nitrate-storing Beggiatoa spp. in coastal marine sediments[END_REF][START_REF] Seitaj | Cable bacteria generate a firewall against euxinia in seasonally hypoxic basins[END_REF]. These are areas of strong upwelling, where high primary productivity and water column stratification promote the formation of oxygen minimum zones (OMZs) in the water column [START_REF] Schmaljohann | Oxygen-minimum zone sediments in the northeastern Arabian Sea off Pakistan: a habitat for the bacterium Thioploca[END_REF]Callbeck et al., 2021). OMZs overlay organicrich, sulfidic sediments characterized by high sulfate reduction rates [START_REF] Gallardo | First Report of the Bacterium Thioploca in the Benguela Ecosystem off Namibia[END_REF][START_REF] Schmaljohann | Oxygen-minimum zone sediments in the northeastern Arabian Sea off Pakistan: a habitat for the bacterium Thioploca[END_REF]Callbeck et al., 2021). In these sediments, CSB form cmthick benthic mats that cover tens of thousands of km 2 of seafloor, from water depths of a few tens of meters to a G. Aloisi et al. / Geochimica et Cosmochimica Acta xxx (2022) [START_REF] Gallardo | Large benthic microbial communities in sulphide biota under Peru-Chile Subsurface Countercurrent[END_REF][START_REF] Fossing | Concentration and Transport of Nitrate by the Mat-Forming Sulfur Bacterium Thioploca[END_REF][START_REF] Jorgensen | Thioploca spp: filamentous sulfur bacteria with nitrate vacuoles[END_REF]. CSB live at the oxic-anoxic interface, where dissolved H 2 S and O 2 or NO 3 -coexist in concentrations smaller than 1 lM, or are separated by a small distance of up to a few cm [START_REF] Teske | The Genera Beggiatoa and Thioploca[END_REF][START_REF] Jørgensen | Big sulfur bacteria[END_REF]. Laboratory experiments with mats collected from natural environments show that CSB adapt to fluctuating chemical gradients, moving vertically though the sediments and storing dissolved NO 3 -and solid S°intracellularly [START_REF] Fossing | Concentration and Transport of Nitrate by the Mat-Forming Sulfur Bacterium Thioploca[END_REF][START_REF] Huettel | Vertical Migration in the Sediment-Dwelling Sulfur Bacteria Thioploca spp. in Overcoming Diffusion Limitations[END_REF][START_REF] Otte | Nitrogen, carbon, and sulfur metabolism in natural Thioploca samples[END_REF][START_REF] Zopfi | Ecology of Thioploca spp.: Nitrate and sulfur storage in relation to chemical microgradients and influence of Thioploca spp. on the sedimentary nitrogen cycle[END_REF][START_REF] Sayama | Impact of bacterial NO 3 transport on sediment biogeochemistry[END_REF][START_REF] Kamp | Anaerobic sulfide oxidation with nitrate by a freshwater Beggiatoa enrichment culture[END_REF][START_REF] Preisler | Biological and chemical sulfide oxidation in a Beggiatoa inhabited marine sediment[END_REF][START_REF] Schutte | Filamentous Giant Beggiatoaceae from the Guaymas Basin Are Capable of both Denitrification and Dissimilatory Nitrate Reduction to Ammonium[END_REF]. When bottom waters are forced to be anoxic, CSB filaments protrude from the sediment, internal stores of NO 3 -are reduced with ambient dissolved H 2 S [START_REF] Fossing | Concentration and Transport of Nitrate by the Mat-Forming Sulfur Bacterium Thioploca[END_REF][START_REF] Otte | Nitrogen, carbon, and sulfur metabolism in natural Thioploca samples[END_REF][START_REF] Sayama | Impact of bacterial NO 3 transport on sediment biogeochemistry[END_REF], and oxidized H 2 S is stored internally as solid S °. Then, when bottom waters are forced to be oxic, CSB filaments retreat into the sediment, and internal S°stores are oxidized to SO 4 2À using bottom water dissolved O 2 or NO 3 -. This metabolic flexibility, together with the vertical motility of filaments, allow CSB to support energy-yielding metabolism for days to months even if electron donors and acceptors are physically separated.

In the field, the use of internal NO 3 -stores to oxidize H 2 S has been observed and quantified by benthic biogeochemical studies of shelf sediments offshore Peru [START_REF] Dale | Biological nitrate transport in sediments on the Peruvian margin mitigates benthic sulfide emissions and drives pelagic N loss during stagnation events[END_REF][START_REF] Sommer | Depletion of oxygen, nitrate and nitrite in the Peruvian oxygen minimum zone cause an imbalance of benthic nitrogen fluxes[END_REF]. To our knowledge, the quantitative field observation of CSB thriving on oxidation of internal S°stores while oxidizing bottom water O 2 or NO 3 -has not been made yet. Based on the interpretation of the fossil filaments as CSB, on the evidence for intense but intermittent biogeochemical activity in marginal basins and on the knowledge of CSB metabolism, in the next section we fully develop the scenario of biogeochemical lowsalinity gypsum formation (Fig. 8, Table 3) that applies principally to the filament-bearing vertically oriented bottom-grown crystals (Fig. 7) of the Primary Lower Gypsum unit.

Low-salinity gypsum precipitation promoted by the biogeochemical S cycle

Our scenario for the involvement of the biogeochemical sulfur cycle in the formation of low-salinity gypsum (Fig. 8; Table 3) relies on the evidence cited above that the deposition of the Primary Lower Gypsum unit was influenced by (i) precession driven climate change, recorded by lithological cycles composed of marls/gypsum couplets (Krijgsman et al., 1999), and (ii) shorter term, seasonal/annual cycles that produced the lamination in the selenite crystals of the gypsum layers (Dela Pierre et al., 2015;Reghizzi et al., 2018). Refer to Table 4 for the stoichiometry of numbered biogeochemical processes in the text that follows.

Marls were deposited during minima of precession, and thus maxima of insolation, when the hydrological cycle is enhanced and marginal basins received the highest fluxes of continental runoff (Lugli et al., 2010;Topper et al., 2014;Dela Pierre et al., 2014;Reghizzi et al., 2018). In these

106CO 2 þ 16N 2 þ H 3 PO 4 þ 154H 2 O ! CH 2 O ðÞ 106 NH 3 ðÞ 16 H 3 PO 4 ðÞ þ 130O 2 þ 16NH 3
At the oxic-anoxic interface, anoxygenic phototrophic bacteria oxidise H 2 St oS O 4 2À (process 2):

H 2 S þ 2CO 2 þ 2H 2 O ! 2CH 2 O þ SO 2À 4 þ 2H þ Deep Stratification
ðÞ þ 138O 2 ! 106CO 2 þ 16HNO 3 þ H 3 PO 4 þ 122H 2 O
Mat when deep water layer is anoxic, NO 3 pumped into mat by Thioploca and reacts with H 2 S to form S°(process 4: DNRA):

HS À þ 0:25NO À 3 þ 1:5CO 2 þ 0:75H 2 O ! S o þ 0:25NH þ 4 þ 1:5HCO À 3
Sulfate reduction produces H 2 S (process 5):

CH 2 O ðÞ 106 NH 3 ðÞ 16 H 3 PO 4 ðÞ þ 53SO 2À 4 ! 67HCO À 3 þ 53HS À þ 122H 2 O Maximum precession = aridity
Arid Surface Convection mixes deep water nutrients into surface water layer NO 3 -based primary production and export to deep layer (process 6):

106CO 2 þ 16HNO 3 þ H 3 PO 4 þ 122H 2 O ! CH 2 O ðÞ
ðÞ þ 138O 2 ! 106CO 2 þ 16HNO 3 þ H 3 PO 4 þ 122H 2 O
Mat Planktic CaCO 3 particles dissolve in mat, releasing Ca 2+ (process 8):

CaCO 3 ! Ca 2þ þ CO 2À 3
Bottom water O 2 pumped by Thioploca into mat from bottom waters, reacts with mat S°producing SO 4 peak (process 9):

S o þ 1:5O 2 þ 2HCO À 3 ! SO 2À 4 þ 2CO 2 þ H 2 O
Seasonal mm-thick gypsum layer precipitates (process 10): -in the surface water was depleted, and N 2 -fixation (diazotrophy, process 1) dominated primary productivity. In deep waters, degradation of settling marine snow (process 3) released nutrients that built-up in the dissolved form. Marine snow that escaped degradation accumulated on the seafloor leading to organic-matter rich marls. Water column stratification also limited the transport of surface dissolved O 2 into deep waters, resulting in decreased deep water O 2 concentrations and, eventually, bottom water anoxia that enhanced organic matter preservation. At the chemocline separating surface oxic from deep anoxic waters, pelagic communities of anoxygenic phototrophic bacteria (process 2) developed. At the benthic boundary, CSB mats thrived by using bottom water NO 3 -to oxidize H 2 S (process 4) produced by sulfate reduction (process 5) in the underlying organic-matter rich, sulfidic sediments, in a process known as dissimilatory nitrate reduction to ammonium (DNRA). While thriving on this catabolic metabolism, CSB stored solid S°intracellularly (process 4).

Ca 2þ þ SO 2À 4 !
Gypsum was deposited at precession maxima (insolation minima) when the average intensity of the hydrological cycle decreased, and the intensity and duration of seasonal aridity was enhanced, driving vertical mixing events in the water column on a yearly basis. Vertical water column mixing transported nutrients from the deep-water column to surface waters, enhancing surface NO 3 --based, possibly diatom-dominated, primary productivity (process 6), while surface O 2 was mixed in bottom waters producing a seasonal event of bottom water oxygenation and aerobic organic matter degradation (process 7). Oxygenated bottom water conditions resulted in a switch of metabolic strategy from DRNA (H 2 S ! S°) to oxidation of S°to SO 4 2À using bottom water O 2 and/or NO 3 (process 9). Inside the diffusion-limited CSB mat, SO 4 2À production resulted in a dissolved SO 4 2À peak, producing gypsum supersaturation and the precipitation of a mm-thick gypsum lamina that rapidly entrapped the biogenic material (filaments, marine snow, diatoms) (process 10). Gypsum precipitation was possibly enhanced by the production of dissolved Ca 2+ within the mat via dissolution of settled planktonic CaCO 3 skeletons (e.g. planktic foraminifers, coccolithophores) (process 8) driven by the acidity produced during S°conversion to SO 4 2À (process 9). This hypothesis is supported by the fact that only siliceous microfossils (diatoms) are preserved in gypsum (Carnevale et al., 2019;Pellegrino et al., 2021), suggesting that the absence of biogenic calcite can reflect a diagenetic bias (Dela Pierre et al., 2014).

During precession maxima, seasonal water column mixing events took place during the arid season, alternated with periods of water column stratification during the humid season, when the marginal basin behaved in a similar way to what happened during precession minima (see above). The rapid rate (hours) at which benthic mats respond to changes in bottom water chemistry in laboratory experiments [START_REF] Fossing | Concentration and Transport of Nitrate by the Mat-Forming Sulfur Bacterium Thioploca[END_REF][START_REF] Huettel | Vertical Migration in the Sediment-Dwelling Sulfur Bacteria Thioploca spp. in Overcoming Diffusion Limitations[END_REF][START_REF] Otte | Nitrogen, carbon, and sulfur metabolism in natural Thioploca samples[END_REF][START_REF] Zopfi | Ecology of Thioploca spp.: Nitrate and sulfur storage in relation to chemical microgradients and influence of Thioploca spp. on the sedimentary nitrogen cycle[END_REF][START_REF] Sayama | Impact of bacterial NO 3 transport on sediment biogeochemistry[END_REF][START_REF] Kamp | Anaerobic sulfide oxidation with nitrate by a freshwater Beggiatoa enrichment culture[END_REF][START_REF] Preisler | Biological and chemical sulfide oxidation in a Beggiatoa inhabited marine sediment[END_REF], makes this hypothesized seasonal shift between S°storage and S°oxidation to SO 4 plausible.

The biogeochemical scenario for low-salinity gypsum formation has to be tested against the d 34 S and d 18 O of gypsum deposits (Fig. 3). More than 80% of the d 34 S SO4 values of dissolved SO 4 2À are included within a very narrow range, from 21 to 23‰ vs CDT (Fig. 3), centered around the Messinian oceanic dissolved sulfate d 34 S composition ($22‰ vs CDT; Masterson et al., 2016). This seems to be at odds with the existence of a dynamic biogeochemical S cycle, because sulfate reduction discriminates strongly against dissolved 34 SO 4 2À [START_REF] Canfield | Biogeochemistry of sulfur isotopes[END_REF][START_REF] Detmers | Diversity of sulfur isotope fractionations by sulfate-reducing prokaryotes[END_REF][START_REF] Bruchert | Controls on stable sulfur isotope fractionation during bacterial sulfate reduction in Arctic sediments[END_REF][START_REF] Habicht | Calibration of sulfate levels in the Archean Ocean[END_REF][START_REF] Canfield | Temperature and its control of isotope fractionation by a sulfate-reducing bacterium[END_REF]Sim et al., 2011), and other processes of the S cycle impart supplementary fractionation to dissolved sulfur species [START_REF] Canfield | Isotope fractionation and sulfur metabolism by pure and enrichment cultures of elemental sulfur-disproportionating bacteria[END_REF][START_REF] Habicht | Sulfur isotope fractionation during bacterial reduction and disproportionation of thiosulfate and sulfite[END_REF][START_REF] Bottcher | Oxygen and sulfur isotope fractionation during anaerobic bacterial disproportionation of elemental sulfur[END_REF][START_REF] Zerkle | Fractionation of multiple sulfur isotopes during phototrophic oxidation of sulfide and elemental sulfur by a green sulfur bacterium[END_REF]Pellerin et al., 2019;[START_REF] Findlay | Sulfide oxidation affects the preservation of sulfur isotope signals[END_REF].

Nevertheless, the invariant d 34 S SO4 could be due to an efficient microbial re-oxidation of H 2 St oS O 4 2À via S°w ithin the CSB mat. The mat exchanges sulfur uniquely via diffusive influx of SO 4 2À from bottom waters and loss of SO 4 2À via gypsum precipitation, while the sink via the precipitation of solid sulfides is negligible (the average reduced S concentration in the Vena del Gesso gypsum, for example, is <170 ppm, data not shown). In these conditions, the d 34 S of the SO 4 2À sink (gypsum precipitation) must be equal to the d 34 S of the SO 42À source (diffusion of SO 4

2À from bottom waters). Evidence exists from a variety of organic-rich sedimentary environments that the efficient anaerobic oxidation of H 2 StoSO 4 2À , known as the ''cryptic sulfur cycle", indeed results in a zone of constant d 34 S SO4 of dissolved SO 4 with sediment depth (Aller et al., 2010;[START_REF] Mills | Geochemical evidence for cryptic sulfur cycling in salt marsh sediments[END_REF][START_REF] Antler | The Sedimentary Carbon-Sulfur-Iron Interplay -A Lesson From East Anglian Salt Marsh Sediments[END_REF].

The d 18 O values of dissolved sulfate deduced from analyses of gypsum (Fig. 3) show a larger range, from 6 to 15‰ vs SMOW. Bacterial sulfate reduction discriminates against 18 O-bearing dissolved SO 4 2À , resulting in an increase in d 18 O of residual sulfate as sulfate reduction proceeds (Brunner et al., 2005;[START_REF] Antler | Coupled sulfur and oxygen isotope insight into bacterial sulfate reduction in the natural environment[END_REF]. In addition to this effect, intracellular oxidation of the sulfite (SO 3 2À ) intermediate to sulfate, and the excretion of the newly produced sulfate to the extracellular domain, introduces oxygen isotopes from water in the dissolved sulfate pool [START_REF] Mizutani | Isotopic behaviour of sulfate oxygen in the bacterial reduction of sulphate[END_REF]Fritz et al., 1989;Wortmann et al., 2007;[START_REF] Farquhar | Sulfur and oxygen isotope study of sulfate reduction in experiments with natural populations from Faellestrand, Denmark[END_REF][START_REF] Turchyn | Kinetic oxygen isotope effects during dissimilatory sulfate reduction: A combined theoretical and experimental approach[END_REF][START_REF] Brunner | The reversibility of dissimilatory sulphate reduction and the cell-internal multi-step reduction of sulphite to sulphide: insights from the oxygen isotope composition of sulphate[END_REF][START_REF] Kohl | No oxygen isotope exchange between water and APS-sulfate at surface temperature: Evidence from quantum chemical modeling and triple-oxygen isotope experiments[END_REF][START_REF] Antler | Coupled sulfur and oxygen isotope insight into bacterial sulfate reduction in the natural environment[END_REF]. Further, oxygen from water molecules is introduced in the dissolved sulfate pool during the re-oxidation of H 2 S [START_REF] Jørgensen | A Thiosulfate Shunt in the Sulfur Cycle of Marine Sediments[END_REF][START_REF] Thamdrup | Thiosulfate and Sulfite Distributions in Porewater of Marine-Sediments Related to Manganese, Iron, and Sulfur Geochemistry[END_REF][START_REF] Canfield | Isotope fractionation and sulfur metabolism by pure and enrichment cultures of elemental sulfur-disproportionating bacteria[END_REF]Canfield et al., , 2005;;[START_REF] Kamyshny | Dynamics of zero-valent sulfur species including polysulfides at seep sites on intertidal sand flats (Wadden Sea, North Sea)[END_REF]. As a result, in sediments characterized by efficient re-oxidation of H 2 St oS O 4 2À , d 18 O SO4 is strongly influenced by the d 18 O of ambient waters (Aller et al., 2010;[START_REF] Mills | Geochemical evidence for cryptic sulfur cycling in salt marsh sediments[END_REF]Blonder et al., 2017).

Considering that the marginal basin where Primary Lower Gypsum and Upper Gypsum deposits formed had a range of d 18 O H2O values (Fig. 2 Given that CSB mats cover vast portions of continental margin sediments overlain by highly-productive and intermittently anoxic water columns [START_REF] Jorgensen | Distribution of colorless sulfur bacteria (Beggiatoa spp.) in a coastal marine sediment[END_REF][START_REF] Gallardo | Large benthic microbial communities in sulphide biota under Peru-Chile Subsurface Countercurrent[END_REF][START_REF] Mussmann | Phylogeny and distribution of nitrate-storing Beggiatoa spp. in coastal marine sediments[END_REF], why doesn't gypsum precipitate in these modern settings? The answer might lie in that these modern systems are less biogeochemically dynamic compared to those that promoted the formation of low salinity gypsum during the Messinian Salinity Crisis. For example, an 11 years long biogeochemical study of water column and benthic biogeochemical processes on the Peru margin [START_REF] Graco | Massive developments of microbial mats following phytoplankton blooms in a naturally eutrophic bay: Implications for nitrogen cycling[END_REF][START_REF] Dale | Strong and Dynamic Benthic-Pelagic Coupling and Feedbacks in a Coastal Upwelling System (Peruvian Shelf)[END_REF] showed that sediments were anoxic throughout this multiannual period, bottom water O 2 varied between 0 and 10 lmol litre À1 , while NO 3 concentrations varied between 0 and 20 lmol litre À1 . In these conditions, benthic CSB communities are forced in the DRNA metabolic mode whereby they use internal NO 3 stores to oxidize sedimentary H 2 St oS 0 . No field evidence of the alternative growth mode (S°! SO 4 2À ) was observed. Further, if anoxic bottom water conditions persist, and degradation of phytodetritus accumulating at the seafloor is intense, then the H 2 S flux from below can outcompete the oxidative ability of CSB, the internal NO 3 stores can be depleted and the CSB mats are destroyed [START_REF] Sommer | Depletion of oxygen, nitrate and nitrite in the Peruvian oxygen minimum zone cause an imbalance of benthic nitrogen fluxes[END_REF]. The Peru mats investigated to date seem to live at the boundary between the DRNA metabolic strategy and being overwhelmed by the benthic H 2 S flux. In summary, the laboratory-observed rapid switching between the two metabolic strategies that CSB are capable of -that is one of the bases of our biogeochemical scenario of low-salinity gypsum formation -is not taking place on the Peru margin.

In summary, we propose that the biogeochemical sulfur cycle did not act alone to produce gypsum saturation, but was one of at least three factors that include the prevailing evaporative conditions and, possibly, Ca 2+ -and SO 4 2À -rich continental runoff. The relative importance of biology and Ca 2+ -and SO 4 2À -rich bearing runoff, as opposed to evaporation, dictated the salinity at which gypsum precipitated.

CONCLUSIONS

In this paper we summarize existing data and present novel geochemical evidence suggesting that gypsum deposits from multiple, geographically separated Mediterranean marginal basins precipitated from low salinity (S < 110 g kg À1 ) water masses. The evidence, which is available primarily for Primary Lower Gypsum deposits, comes from (i) the unusually low salinities (mostly < 50 g kg À1 )in gypsum fluid inclusions from the Sorbas basin (Spain), the Catanzaro Trough (southern Italy), the Vena del Gesso basin (northern Italy) and the Piedmont basin (northern Italy); and (ii) unusually light oxygen isotope composition (d 18 O<0 ‰ vs SMOW) of gypsum parent waters from the Piedmont basin (northern Italy) and the Sorbas basin (Spain). We propose that multiple factors, including evaporative conditions, Ca 2+ -and SO 4 2À -bearing runoff and the biogeochemical sulfur cycle may have acted simultaneously to variable degrees, leading to a span of salinities at gypsum saturation from the thermodynamic value of 110 g kg À1 ,to salinities lower that modern seawater (<35 g kg À1 ). The involvement of biogeochemistry, via the production of a microbially-mediated, and temporally transient, increase of dissolved sulfate in a CSB-dominated benthic microbial mat, is suggested by independent petrographic and organic geochemical evidence. Should our scenario for the formation of low-salinity gypsum be confirmed for other geological epochs, it would imply that large gypsum deposits do not require evaporation to increases seawater salinity in excess of 110 g kg À1 , expanding the range of environments that promote marine gypsum deposition. It would also imply that an additional, biological coupling between the calcium, sulfur and carbon cycles exists. The late Miocene evolution of the Mediterranean Basin is characterized by major changes in connectivity, climate and tectonic activity resulting in unprecedented environmental and ecological disruptions. During the Messinian Salinity Crisis (MSC, this culminated in most scenarios first in the precipitation of gypsum around the Mediterranean margins (Stage 1, 5.97-5.60 Ma) and subsequently > 2 km of halite on the basin floor, which formed the so-called Mediterranean Salt Giant (Stage 2,. The final MSC Stage 3, however, was characterized by a "low-salinity crisis", when a second calcium-sulfate unit (Upper Gypsum; substage 3.1, showing (bio)geochemical evidence of substantial brine dilution and brackish biota-bearing terrigenous sediments (substage 3.2 or Lago-Mare phase, 5.42-5.33 Ma) deposited in a Mediterranean that received relatively large amounts of riverine and Paratethys-derived low-salinity waters. The transition from hypersaline evaporitic (halite) to brackish facies implies a major change in the Mediterranean's hydrological regime. However, even after nearly 50 years of research, causes and modalities are poorly understood and the original scientific debate between a largely isolated and (partly) desiccated Mediterranean or a fully connected and filled basin is still vibrant. Here we present a comprehensive overview that brings together (chrono)stratigraphic, sedimentological, paleontological, geochemical and seismic data from all over the Mediterranean. We summarize the paleoenvironmental, paleohydrological and paleoconnectivity scenarios that arose from this cross-disciplinary dataset and we discuss arguments in favour of and against each scenario.

Introduction

At the end of the Miocene, orbital and tectonic drivers combined to alter the amount of water delivered to the Mediterranean Basin by the Atlantic Ocean from the west, the brackish Eastern Paratethys (i.e. Euxinic-Caspian Basin system) from the east and the major peri-Mediterranean freshwater drainage systems (e.g. African rivers and Rhône; [START_REF] Griffin | Aridity and humidity: two aspects of the late Miocene climate of North Africa and the Mediterranean[END_REF][START_REF] Gladstone | The Mediterranean hydrologic budget from a Late Miocene global climate simulation[END_REF]Van der Laan et al., 2006;[START_REF] Hilgen | Astronomical tuning as the basis for high resolution chronostratigraphy: the intricate history of the Messinian Salinity Crisis[END_REF]Ryan, 2009;Flecker et al., 2015;[START_REF] Marzocchi | Orbital control on late Miocene climate and the North African monsoon: insight from an ensemble of sub-precessional simulations[END_REF]Marzocchi et al., , 2016[START_REF] Marzocchi | Precessional drivers of late Miocene Mediterranean sedimentary sequences: African summer monsoon and Atlantic winter storm tracks[END_REF]Simon et al., 2017;Krijgsman et al., 2018;[START_REF] Capella | Mantle resistance against Gibraltar slab dragging as a key cause of the Messinian Salinity Crisis[END_REF]. The changes in extra and intrabasinal connectivity resulted in unprecedented paleoceanographic and paleohydrological budget changes that led to a relatively short-lived environmental and ecological crisis (approx. 660 kyr; 5.97-5.33 Ma), for which the term Messinian Salinity Crisis (MSC) was coined [START_REF] Selli | Il Bacino del Metauro[END_REF][START_REF] Selli | Il Messiniano Mayer-Eymar 1867. Proposta di un neostratotipo[END_REF]. Most conspicuous was the rapid accumulation of several kilometers of halite (i.e. ~1 million km 3 ) on the Mediterranean abyssal plains (e.g. [START_REF] Hsü | Origin of Saline Giants: a critical review after the discovery of the Mediterranean evaporite[END_REF]Ryan, 1973;[START_REF] Montadert | Messinian event: seismic evidence[END_REF]Haq et al., 2020). This happened within 50 kyr, from 5.60-5.55 Ma, according to Roveri et al. (2014a) and [START_REF] Manzi | The onset of the Messinian salinity crisis in the deep Eastern Mediterranean basin[END_REF], or in >300 kyr, when starting at 5.97 Ma, as put forward by [START_REF] Meilijson | Deep-basin evidence resolves a 50-year-old debate and demonstrates synchronous onset of Messinian evaporite deposition in a non-desiccated Mediterranean[END_REF]Meilijson et al. ( , 2019)).

During the ~200 kyr lapse (i.e. MSC Stage 3 following Roveri et al., 2014a;Fig. 1a) between the end of salt precipitation (5.55 Ma) and the restoration of the still enduring marine conditions (5.33 Ma), the Mediterranean underwent a sequence of paleohydrological and base-level changes that are the topic of intense and long-standing debates. The initial and still widely endorsed hypothesis was that the Mediterranean Sea, following the major drawdown event that led to halite deposition (i. e. Stage 2), maintained the isolated, deeply-desiccated geography containing a series of hypersaline (substage 3.1; 5.55-5.42 Ma) and hyposaline (substage 3.2; 5.42-5.33 Ma) ponds which only received water from local streams and were colonized by Black Sea organisms carried by aquatic migratory birds (Fig. 1b; e.g. Ruggieri, 1967;[START_REF] Decima | Comments on late Messinian microfaunas in several sections from Sicily[END_REF][START_REF] Decima | Ostracodi del Gen. Cyprideis JONES del Neogene e del Quaternario italiani[END_REF]Wezel, 1971, 1973;Cita et al., 1978;Müller et al., 1990;Benson and Rakic-El Bied, 1991;Benson et al., 1991;[START_REF] Müller | Origin and age of the Mediterranean Messinian evaporites: implications from Sr isotopes[END_REF][START_REF] Butler | Tectonics and sequence stratigraphy in Messinian Basins, Sicily: constraints on the initiation and termination of the Mediterranean salinity crisis[END_REF][START_REF] Orszag-Sperber | La transition Messinien-Pliocène en Méditerranée orientale (Chypre): la période du Lago-Mare et sa signification[END_REF][START_REF] Rouchy | Paleoenvironmental changes at the Messinian-Pliocene boundary in the eastern Mediterranean (southern Cyprus basins): significance of the Messinian Lago-Mare[END_REF][START_REF] Kartveit | Evidence of sea level drawdown at the end of the Messinian salinity crisis and seismic investigation of the Nahr Menashe unit in the northern Levant Basin, offshore Lebanon[END_REF][START_REF] Madof | Discovery of vast fluvial deposits provides evidence for drawdown during the late Miocene Messinian salinity crisis[END_REF][START_REF] Camerlenghi | Seismic markers of the Messinian salinity crisis in the deep Ionian Basin[END_REF][START_REF] Caruso | The late Messinian "Lago-Mare" event and the Zanclean Reflooding in the Mediterranean Sea: new insights from the Cuevas del Almanzora section (Vera Basin, South-Eastern Spain)[END_REF]Raad et al., 2021). As morphological and seismic reflection studies at the Strait of Gibraltar documented a ~400 km long erosional trough connecting the Gulf of Cadiz (Atlantic Ocean) to the Mediterranean Sea, this scenario of a lowered Mediterranean Sea was promptly linked to the termination of the MSC [START_REF] Mckenzie | From desert to deluge in the Mediterranean[END_REF][START_REF] Blanc | The opening of the Plio-Quaternary Gibraltar Strait: assessing the size of a cataclysm[END_REF][START_REF] Garcia-Castellanos | Catastrophic flood of the Mediterranean after the Messinian salinity crisis[END_REF], 2020). This conclusion has recently been reinforced by the discovery of vast chaotic deposits sitting at the claimed Miocene/Pliocene transition in the area of the Malta Escarpment-Ionian Abyssal Plain [START_REF] Micallef | Evidence of the Zanclean megaflood in the eastern Mediterranean Basin[END_REF][START_REF] Micallef | Geomorphic evolution of the Malta Escarpment and implications for the Messinian evaporative drawdown in the eastern Mediterranean Sea[END_REF][START_REF] Spatola | A single-stage megaflood at the termination of the Messinian salinity crisis: geophysical and modelling evidence from the eastern Mediterranean Basin[END_REF].

In more recent years, the desiccated basin model was challenged by the observation of deposits that are uniform in terms of sedimentology and stratigraphic architecture (Roveri et al., 2008a), ostracod content [START_REF] Gliozzi | Paratethyan ostracod immigrants in Italy during the Late Miocene[END_REF]Stoica et al., 2016) and geochemistry [START_REF] Mcculloch | Sr isotope constraints on the Mediterranean environment at the end of the Messinian salinity crisis[END_REF]García-Veigas et al., 2018;Andreetto et al., 2021) throughout the Mediterranean marginal belt and of δD n-alkanes and δD alkenones sharing similarities with the coeval Atlantic Ocean and Black Sea, respectively (Vasiliev et al., 2017). A model of a (relatively) full Mediterranean Sea developed (Fig. 1c), where the debate mainly concerns the provenance of the hydrological fluxes and the resultant hydrochemical composition of the water mass. In this scenario, the Mediterranean was first, during substage 3.1, transformed into a new gypsum-precipitating basin filled with marine and continent-derived (Roveri et al., 2014a). Stage 3, here of interest, spans between 5.55 Ma and 5.332 Ma, the astronomical ages of the base of the Upper Gypsum Unit (following [START_REF] Manzi | A new facies model for the Upper Gypsum of Sicily (Italy): chronological and paleoenvironmental constraits for the Messinian salinity crisis in the Mediterranean[END_REF]Trubi Formation (Van Couvering et al., 2000) waters (e.g. [START_REF] Manzi | A new facies model for the Upper Gypsum of Sicily (Italy): chronological and paleoenvironmental constraits for the Messinian salinity crisis in the Mediterranean[END_REF]Roveri et al., 2014c;Flecker et al., 2015;Vasiliev et al., 2017;García-Veigas et al., 2018;Grothe et al., 2020). Then, during substage 3.2, it became a brackish lake-sea comparable to the present-day Black Sea or Caspian Sea (Roveri et al., 2008a;Stoica et al., 2016;Andreetto et al., 2021), depending on whether a marine connection with the Atlantic was active [START_REF] Manzi | A new facies model for the Upper Gypsum of Sicily (Italy): chronological and paleoenvironmental constraits for the Messinian salinity crisis in the Mediterranean[END_REF]Roveri et al., 2014bRoveri et al., , 2014c;;Flecker et al., 2015;Marzocchi et al., 2016;Vasiliev et al., 2017;García-Veigas et al., 2018) or not (e.g. [START_REF] Mcculloch | Sr isotope constraints on the Mediterranean environment at the end of the Messinian salinity crisis[END_REF]Roveri et al., 2008a), and with a base-level fluctuating by hundreds of meters with precessional periodicity [START_REF] Fortuin | The Messinian of the Nijar basin (SE Spain): sedimentation, depositional environments and paleogeographic evolution[END_REF][START_REF] Ben Moshe | Estimating drawdown magnitudes of the Mediterranean Sea in the Levant basin during the Lago Mare stage of the Messinian Salinity Crisis[END_REF]Andreetto et al., 2021). In the relatively full scenario, the revival of marine conditions is ascribed to either connectivity changes (Marzocchi et al., 2016) or to a moderate sea-level rise (Andreetto et al., 2021). In contrast, Carnevale et al. (2006aCarnevale et al. ( , 2006b[START_REF] Carnevale | Did the Mediterranean marine reflooding precede the Miocene/Pliocene boundary? Paleontological and geochemical evidence from upper Messinian sequences of Tuscany, Italy. Paleogeogr[END_REF][START_REF] Carnevale | Fossil marine fishes and the "Lago Mare" event: Has the Mediterranean ever transformed into a brackish lake?[END_REF] and [START_REF] Grunert | Estuarine Lago Mare fauna from the Tertiary Piedmont Basin indicates episodic Atlantic/Mediterranean exchange during the final stage of the Mediterranean Salinity Crisis[END_REF], based on the recovery of fish remains ascribed to marine species, proposed that fully marine conditions were in force in the Mediterranean already at the end of substage 3.1.

After nearly 50 years of research on both onshore and offshore localities (Fig. 2), the observations backing up the competing desiccated and full-basin Mediterranean models remain extremely difficult to reconcile. Uncertainties regarding the chronostratigraphic framework of Stage 3 deposits, the origin and migration of its characteristic biota, the meaning of the data derived from the applied geochemical techniques [START_REF] Roveri | The Terminal Carbonate Complex: the record of sea-level changes during the Messinian salinity crisis[END_REF](Roveri et al., , 2019a)); 4-Nijar Basin [START_REF] Fortuin | The Messinian of the Nijar basin (SE Spain): sedimentation, depositional environments and paleogeographic evolution[END_REF]; 5-Vera Basin [START_REF] Fortuin | The enigmatic Messinian -Pliocene section of Cuevas del Almanzora (Vera Basin, SE Spain) revisited-erosional features and strontium isotope ages[END_REF]; 6-Bajo Segura Basin [START_REF] Soria | The stratigraphic record of the Messinian salinity crisis in the northern margin of the Bajo Segura basin (SE Spain)[END_REF](Soria et al., , 2008a(Soria et al., , 2008b)); 7-Mallorca (Mas and Fornós, 2020); 8-Melilla Basin [START_REF] Rouchy | Late Messinian to Early Pliocene paleoenvironmental changes in the Melilla Basin (NE Morocco) and their relation to Mediterranean evolution[END_REF]; 9-Boudinar Basin [START_REF] Merzeraud | Sedimentology and sequence stratigraphy of the late-Messinian-Early pliocene continental to marine deposits of the Boudinar basin (North Morocco)[END_REF]; 10-Chelif Basin [START_REF] Rouchy | The end of the Messinian salinity crisis: evidences from the Chelif Basin (Algeria)[END_REF]; 11-Sahel area [START_REF] Frigui | Evidences of "Lago-Mare" episode around the Messinian-Pliocene boundary in eastern Tunisia (central Mediterranean)[END_REF]; 12-Aléria Basin and 13-Rhône Valley [START_REF] Carbonnel | La zone à Loxoconcha djaffarovi Schneider (ostracodes, Miocène supérieur) ou le Messinien de la vallée du Rhône[END_REF]; 14-Piedmont Basin (Dela Pierre et al., 2011, 2016); 15-Po Plain [START_REF] Ghielmi | Sedimentary and tectonic evolution in the eastern Po-Plain and northern Adriatic Sea area from Messinian to Middle Pleistocene (Italy)[END_REF][START_REF] Ghielmi | Late Miocene-Middle Pleistocene sequences in the Po Plain-Northern Adriatic Sea (Italy): the stratigraphic record of modification phases affecting a complex foreland basin[END_REF][START_REF] Amadori | Restored topography of the Po Plain-Northern Adriatic region during the Messinian base level drop-Implications for the physiography and compartmentalization of the paleo-Mediterranean basin[END_REF]; 16-Fine Basin (Cava Serredi section; Carnevale et al., 2006a[START_REF] Carnevale | Did the Mediterranean marine reflooding precede the Miocene/Pliocene boundary? Paleontological and geochemical evidence from upper Messinian sequences of Tuscany, Italy. Paleogeogr[END_REF]. 17-21 Apennine system: Romagna sections (17, [START_REF] Roveri | Stratigraphy of the Messinian post-evaporitic stage in eastern-Romagna (northern Apennines, Italy)[END_REF], Trave section (18, [START_REF] Iaccarino | The Trave section (Monte dei Corvi, Ancona, central Italy): an integrated paleontological study of the Messinian deposits[END_REF], Maccarone section (19, [START_REF] Bertini | The Northern Apennines palynological record as a contribute for the reconstruction of the Messinian paleoenvironments[END_REF], Grossi et al., 2008;[START_REF] Sampalmieri | Paleoenvironments of the Mediterranean Basin at the Messinian hypersaline/ hyposaline transition: evidence from natural radioactivity and microfacies of postevaporitic successions of the Adriatic sub-basin[END_REF]Pellen et al., 2017), Colle di Votta (20)-Fonte dei Pulcini (21)-Stingeti ( 22) sections [START_REF] Cosentino | High-frequency cyclicity in the latest Messinian Adriatic foreland basin: insight into paleoclimate and paleoenvironments of the Mediterranean Lago-Mare episode[END_REF][START_REF] Cosentino | Orbitally forced paleoenvironmental and paleoclimate changes in the late postevaporitic Messinian of the central Mediterranean Basin[END_REF][START_REF] Cosentino | Refining the Mediterranean "Messinian gap" with high-precision U-Pb zircon geochronology, central and northern Italy[END_REF](Cosentino et al., , 2018)), Mondragone 1 well (23, Cosentino et al., 2006), Crotone Basin (24, Roveri et al., 2008a); 25-27 Sicily: Villafranca Tirrena (25) and Licodia Eubea ( 26) sections [START_REF] Sciuto | New reports of Paratethyan ostracods affinity from the Mediterranean Basin (Sicily, Italy)[END_REF]), Caltanissetta Basin (27, Manzi et al., 2009); 28-Corfu [START_REF] Pierre | Reconstruction of the paleoenvironmental changes around the Miocene-Pliocene boundary along a West-East transect across the Mediterranean[END_REF]; 29-Zakinthos (Karakitsios et al., 2017b); 30-Crete [START_REF] Cosentino | The late Messinian Lago-Mare episode in the Mediterranean Basin: preliminary report on the occurrence of Paratethyan ostracod fauna from central Crete (Greece)[END_REF]; 31-Cyprus [START_REF] Rouchy | Paleoenvironmental changes at the Messinian-Pliocene boundary in the eastern Mediterranean (southern Cyprus basins): significance of the Messinian Lago-Mare[END_REF]Manzi et al., 2016a); 32-Adana Basin [START_REF] Radeff | Stratigraphic architecture of the upper Messinian deposits of the Adana Basin (southern Turkey): implications for the Messinian salinity crisis and the Taurus petroleum system[END_REF].

F. Andreetto et al. and the relationship between the Mediterranean and its surrounding water bodies (i.e. Atlantic Ocean, Indian Ocean and Paratethys) all inhibit a clear understanding of the Mediterranean base-level and its hydrochemical structure.

In this paper we attempt to summarize all the existing, but heavily scattered, data resulting from ~50 years of cross-disciplinary studies with the aim of providing a comprehensive overview of the stratigraphic arrangement of Stage 3 onshore and offshore deposits, as well as of their sedimentological, paleontological, geochemical and seismic properties. Subsequently, we assemble the observations favoring both end-member scenarios of a relatively desiccated and relatively full Mediterranean. Finally, we focus on novel future analytical techniques and approaches that have the potential to constrain Mediterranean base-level during MSC Stage 3 as well as the changing hydrological fluxes and connectivity phases between the intra-Mediterranean basins and the neighboring Atlantic Ocean and Paratethyan domains as a mean of reconstructing the state of the art of the complex history of this enigmatic period of the Mediterranean history once and for all.

The terminal Stage 3 of the MSC

Historic overview of nomenclature and concepts

The final phase of the MSC (i.e. substage 3.2), also known as "Lago-Mare", finds its sedimentary expression in cyclically-arranged terrigenous and evaporitic sediments hosting unique faunal assemblages of ostracods, mollusks and dinoflagellate cysts (dinocysts). They are related, at species level, to those inhabiting, during the Miocene, the brackish basins of the Paratethys realm (e.g. [START_REF] Gliozzi | Paratethyan ostracod immigrants in Italy during the Late Miocene[END_REF]Stoica et al., 2016). But what exactly is the "Lago-Mare"? This widely employed expression in the MSC literature encompasses a variety of meanings that make its application doubtful and misleading. The root of the wording "Lago-Mare" is to be found in the Russian scientific literature of the late 1800s. Nikolai Andrusov (1890) used the corresponding Russian term with a geographical and chronological connotation in reference to the series of central-eastern European basins that during the Miocene turned from marine settings to desalinized semi-isolated lakes with an endemic fresh-brackish water biota association (e.g. Popov et al., 2006 and references therein). The original monograph of [START_REF] Andrusov | Les Dreissenidae fossiles et actuelles d'Eurasie[END_REF] was not widely available outside Russia, but his attendance of international conferences allowed his research to spread outside the Russian borders. From the publications of the French geologists Suzette Gillet [START_REF] Gillet | Essai de classification du Miocène superieur et du Pliocene inferieur de Roumanie. La Transylvanie et le Banat[END_REF][START_REF] Gillet | Essai de classification du Miòcene supérieur et du Pliocène inférieur de Roumanie. Le bassin dacique[END_REF] and Maurice Gignoux (Gignoux, 1936a) we can state with relative certainty that the original meaning of the word "Lago-Mare" (here reported with the French counterpart "Lac-Mer") had its provenance in the Russian literature: [START_REF] Gillet | Essai de classification du Miocène superieur et du Pliocene inferieur de Roumanie. La Transylvanie et le Banat[END_REF]. " […] During the Volhynien (Sarmatique inferior) there was a lac-mer of uniform fauna that extended through all the eastern Europe. […] and the fauna of the eastern regions of that huge lac-mer was completely differenciated [sic] from the one in the western regions. […]" [START_REF] Gillet | Essai de classification du Miòcene supérieur et du Pliocène inférieur de Roumanie. Le bassin dacique[END_REF].

"[…] An isolation of the basin, that became a brackish, isolated basin. Then, a uniform fauna populated this immense lac-mer which was divided […] into Pannonian basin, […] Dacique Basin, and Euxin and Caspian basin […]"

"[…]

The Pontien fauna is not anymore a fauna characteristic of an internal saline sea, as in the Sarmantien, but is a fauna of a "desalinated lagoon", a lac-mer, as the Russian geologists named it. […]" (Gignoux, 1936b).

In the late 19 th [START_REF] Capellini | Gli strati a Congerie o la formazione gessoso solfifera nella provincia di Pisa e nei dintorni di Livorno[END_REF] and 20 th century [START_REF] Ogniben | Le argille scagliose del Crotonese[END_REF][START_REF] Ruggieri | La serie marine pliocenica e quaternaria della Val Marecchia: Atti Accad[END_REF]Ruggieri, , 1967;;[START_REF] Decima | Ostracodi del Gen. Cyprideis JONES del Neogene e del Quaternario italiani[END_REF], late Messinian ostracod-and mollusk-bearing deposits in the Mediterranean were described at several Italian localities. Initially, the expressions "Congeria beds" [START_REF] Capellini | Gli strati a Congerie o la formazione gessoso solfifera nella provincia di Pisa e nei dintorni di Livorno[END_REF] and "Melanopsis beds" [START_REF] Ruggieri | La serie marine pliocenica e quaternaria della Val Marecchia: Atti Accad[END_REF] were used. Later on, Ruggieri (1967) pointed out the affinity of these faunal elements with those of the Pontian of the Paratethys. Consequently, he speculated on a feasible Paratethys-like paleoenvironmental configuration for the Mediterranean in the latest Messinian and he coined the Italian translation (i. e. "Lago-Mare") from the French "Lac-Mer" in reference to the shallowwater lakes claimed to be widely distributed across the Mediterranean. Progress in the 1970s in onshore and offshore exploration highlighted the temporally well-constrained distribution of the Paratethyan organisms in the Mediterranean [START_REF] Carbonnel | La zone à Loxoconcha djaffarovi Schneider (ostracodes, Miocène supérieur) ou le Messinien de la vallée du Rhône[END_REF]. On this premise, Hsü et al. (1978a) proposed to use "Lago-Mare" to "designate the latest Messinian oligohaline environment, postdating evaporite deposition and predating Pliocene marine sedimentation […] in order to distinguish it from "lac mer" which, strictly speaking, was a Paratethyan environment". Notwithstanding the new definition, in various parts of the text they used "Lago-Mare" to refer to the Paratethyan lakes (pp. 1071-1072: "[…] The upper Messinian Mediterranean was floored by a series of desert basins, some with salt lakes, prior to inundation by the Lago-Mare."), thus giving rise to the confusion on how to use the term properly.

In the most recent stratigraphic overview of the MSC (Fig. 1a; Roveri et al., 2014a), the terminal MSC stage is called Stage 3, which is in turn subdivided into substages 3.1 and 3.2 (also termed Lago-Mare). Beside such a chronostratigraphic definition, the term "Lago-Mare" has also been used for a typical biofacies of the late Messinian Mediterranean (e. g. [START_REF] Fortuin | The enigmatic Messinian -Pliocene section of Cuevas del Almanzora (Vera Basin, SE Spain) revisited-erosional features and strontium isotope ages[END_REF][START_REF] Gliozzi | A late Messinian brackish water ostracod fauna of Paratethyan aspect from Le Vicenne Basin (Abruzzi, central Apennines, Italy)[END_REF][START_REF] Gliozzi | Late Messinian lago-mare ostracod paleoecology: a correspondence analysis aroach[END_REF][START_REF] Sciuto | New reports of Paratethyan ostracods affinity from the Mediterranean Basin (Sicily, Italy)[END_REF], for the pelitic beds encasing the Paratethyan-related fauna (i.e. a lithofacies; e.g. [START_REF] Fortuin | The Messinian of the Nijar basin (SE Spain): sedimentation, depositional environments and paleogeographic evolution[END_REF][START_REF] Sciuto | New reports of Paratethyan ostracods affinity from the Mediterranean Basin (Sicily, Italy)[END_REF], as the name of an informal lithostratigraphic unit (usually distinguished by its fossil content) sandwiched between the Sicilian Upper Gypsum and the Arenazzolo Fm. (Fig. 4b; [START_REF] Clauzon | Influence of Mediterranean sea-level changes on the Dacic Basin (Eastern Paratethys) during the late Neogene: the Mediterranean Lago Mare facies deciphered[END_REF][START_REF] Londeix | Messinian paleoenvironments and hydrology in Sicily (Italy): the dinoflagellate cyst record[END_REF][START_REF] Popescu | Galeacysta etrusca complex: dinoflagellate cyst marker of Paratethyan influxes to the Mediterranean Sea before and after the peak of the Messinian Salinity Crisis[END_REF][START_REF] Bache | A twostep process for the reflooding of the Mediterranean after the Messinian Salinity Crisis[END_REF] and to denote multiple (3 to 4) spilling events of the Paratethys into the Mediterranean [START_REF] Clauzon | Influence of Mediterranean sea-level changes on the Dacic Basin (Eastern Paratethys) during the late Neogene: the Mediterranean Lago Mare facies deciphered[END_REF][START_REF] Clauzon | New insights on the Sorbas Basin (SE Spain): the onshore reference of the Messinian salinity crisis[END_REF]Popescu et al., 2007[START_REF] Popescu | Galeacysta etrusca complex: dinoflagellate cyst marker of Paratethyan influxes to the Mediterranean Sea before and after the peak of the Messinian Salinity Crisis[END_REF][START_REF] Popescu | Lago Mare episodes around the Messinian-Zanclean boundary in the deep southwestern Mediterranean[END_REF][START_REF] Suc | The Messinian salinity crisis in the Dacic Basin (SW Romania) and early Zanclean Mediterranean-Paratethys high sea-level connection[END_REF][START_REF] Bache | A twostep process for the reflooding of the Mediterranean after the Messinian Salinity Crisis[END_REF][START_REF] Couto | Lago Mare and the Messinian salinity crisis: evidences from the Alboran Sea (S. Spain)[END_REF][START_REF] Frigui | Evidences of "Lago-Mare" episode around the Messinian-Pliocene boundary in eastern Tunisia (central Mediterranean)[END_REF][START_REF] Mas | The messinian salinity crisis in Mallorca: New insights for a western mediterranean stratigraphic scenario[END_REF].

This being a review, we use the widely employed definition of the model of Roveri et al., 2014a) (Fig. 1a) and regard the Lago-Mare as a "phase of massive biota migration from the Paratethys realm, cyclostratigraphically constrained between 5.42 Ma and 5.332 Ma (Roveri et al., 2008a;Grossi et al., 2011), during which the Mediterranean sedimentary environments underwent an impressive freshening". Nevertheless, we call for caution in the use of this definition of "Lago-Mare" in future studies, since 5.42 Ma as the (astronomical) age of the first entrance of Paratethyan organisms into the Mediterranean is likely to be incorrect (see subsection 5.1) and evidence of 'impressive freshening' are already present much earlier (e.g. at Eraclea Minoa; Vasiliev et al., 2017;García-Veigas et al., 2018).

Development of a chronostratigraphic framework

Issues of the timing and duration of the MSC only began to be tackled in the 1990s, in parallel with discussion concerning the nature of its extreme paleoenvironments [START_REF] Schmalz | Deep-water evaporite deposition, a genetic model[END_REF]Hsü et al., 1973a, 1973b, 1973c, Hsü et al., 1978a[START_REF] Hsü | Initial reports of the deep sea drilling project: DSDP volume XLII Part[END_REF][START_REF] Nesteroff | Un modéle pour les évaporites messinienne en Méditerranée, bassins peu profonds avec dépot d'évaporites lagunaires[END_REF][START_REF] De Benedetti | The problem of the origin of the salt deposits in the Mediterranean and of their relations to other salt occurrences in the Neogene formations of the contiguous regions[END_REF]. While published models [START_REF] Butler | Tectonics and sequence stratigraphy in Messinian Basins, Sicily: constraints on the initiation and termination of the Mediterranean salinity crisis[END_REF][START_REF] Andreetto | Alternate interpretation of the Messinian salinity crisis, controversy resolved?[END_REF]Krijgsman et al., 1999a;Rouchy and Caruso, 2006) mostly converged on the (astronomical) age of the marine replenishment at the beginning of the Pliocene (5.332 Ma;[START_REF] Van Couvering | The base of the Zanclean Stage and of the Pliocene Series[END_REF], there were disagreements about the age of the onset of the MSC (synchronous vs diachronous) and of specific events within it (see discussion in Roveri et al., 2014a). Among these, the work of Krijgsman et al. (1999a) has obtained wide consensus. Their cyclostratigraphic tuning and correlation of continuous and bio-magnetostratigraphically constrained preevaporitic sections in Spain (Sorbas), Sicily (Gibliscemi/Falconara) and Greece (Metochia) resulted in a synchronous age of 5.96±0.02 Ma for the MSC onset (later refined to 5.97 Ma by Manzi et al., 2013). The astronomical ages for the onset (Krijgsman et al., 1999a) and termination [START_REF] Van Couvering | The base of the Zanclean Stage and of the Pliocene Series[END_REF] of the MSC are not contentious since F. Andreetto et al. the characteristic sedimentary cyclicity and sediments' properties (e.g. color of the lithologies and biota content) of the pre-and post-MSC successions fit robustly with the insolation curve (see also Van der Laan et al., 2006 andTopper andMeijer, 2015).

The cyclic arrangement of the MSC sediments (Fig. 3a) led scientists to interpret that the same cyclostratigraphic approach could be used to gain precise dates for events within the MSC (e.g. [START_REF] Hilgen | Extending the astronomical (polarity) time scale into the Miocene[END_REF][START_REF] Vai | Chapter E3 Cyclostratigraphic estimate of the messinian stage duration[END_REF]Krijgsman et al., 1999b[START_REF] Krijgsman | Astrochronology for the Messinian Sorbas basin (SE Spain) and orbital (precessional) forcing for evaporite cyclicity[END_REF], bypassing the challenge posed by the unsuitability of the classic biomagnetostratigraphic tools for the MSC successions. Characteristic interference patterns of eccentricity and precession have been tentatively recognized in the Sicilian Eraclea Minoa section (see subsection 3.8;Van der Laan et al., 2006). However, clear orbital signals are typically poorly expressed in MSC records and, when they are present, like in Sicily, they are not (vertically) repeated with sufficient frequency to establish clear phase relations with the astronomical cyclicity. For this reason, the simple counting of cycles with no analysis of cyclostratigraphic pattern in proxy records has mostly been employed as a correlation method (Roveri et al., 2008a;[START_REF] Manzi | A new facies model for the Upper Gypsum of Sicily (Italy): chronological and paleoenvironmental constraits for the Messinian salinity crisis in the Mediterranean[END_REF]Manzi et al., 2016a;[START_REF] Cosentino | Refining the Mediterranean "Messinian gap" with high-precision U-Pb zircon geochronology, central and northern Italy[END_REF].

The age of the base of Stage 3 is largely determined by correlating the sedimentary cycles of the Upper Gypsum unit (UG) at Eraclea Minoa (Sicily) with the astronomical curve La2004 [START_REF] Laskar | A long term numerical solution for the insolation quantities of the Earth[END_REF]. The UG sedimentary cyclicity consists of alternating gypsum and mudstone beds of variable thickness (Figs. 5g-i; see subsection 3.8). Precessional variation of the Mediterranean freshwater budget tied tightly to the African monsoon and Atlantic storms are the drivers interpreted to lie behind the gypsum-mudstone cycles (e.g. [START_REF] Marzocchi | Orbital control on late Miocene climate and the North African monsoon: insight from an ensemble of sub-precessional simulations[END_REF][START_REF] Marzocchi | Precessional drivers of late Miocene Mediterranean sedimentary sequences: African summer monsoon and Atlantic winter storm tracks[END_REF]Simon et al., 2017). Variations of the freshwater discharge cause the pycnocline to shift vertically, resulting in brine concentration and gypsum precipitation during to the arid/dry phases of the precession cycles (precession maxima-insolation minima) and brine dilution and mudstone deposition during the humid/wet phases (precession minimainsolation maxima) ( Van der Laan et al., 2006;[START_REF] Manzi | A new facies model for the Upper Gypsum of Sicily (Italy): chronological and paleoenvironmental constraits for the Messinian salinity crisis in the Mediterranean[END_REF]. Two different tuning options exist in literature ( Van der Laan et al., versus Manzi et al., 2009;Fig. 3a):

1. Van der Laan et al. ( 2006) tentatively recognized sedimentary patterns that they correlated with the astronomical curves by using the same phase relationships between the sedimentary cycles and the astronomical cycles as are seen in Plio-Pleistocene sapropel-bearing marine successions of the Mediterranean [START_REF] Hilgen | Astronomical calibration of Gauss to Matuyama sapropels in the Mediterranean and implication for the geomagnetic polarity time scale[END_REF]. The four closely spaced gypsum beds III to VI were regarded as a cluster, i.e. the sedimentary expression of a 100 kyr eccentricity maximum [START_REF] Hilgen | Astronomical calibration of Gauss to Matuyama sapropels in the Mediterranean and implication for the geomagnetic polarity time scale[END_REF][START_REF] Strasser | Cyclostratigraphy-concepts, definitions, and alications[END_REF], whereas the preceding and following evaporite-free marly interval were attributed to a phase of low-amplitude precession oscillations caused by a 100 kyr eccentricity minimum (Fig. 3a). Tuning downward from the conformable Miocene/Pliocene boundary (Fig. 6d) and arguing that the precession peak at ~5.38 Ma, which has an extremely low amplitude, is not expressed in the sedimentary record, Van der Laan et al. ( 2006) correlated gypsums III to VI with the four successive precession/ insolation peaks of the 100 kyr eccentricity maximum dated around 5.44 Ma and the overlying and underlying gypsum-free marly interval fell within 100 kyr eccentricity minimum cycles (Fig. 3a, right log). This tuning resulted in an astronomical age of ~5.51 Ma for the first gypsum bed in their log (i.e. gypsum II in the log of [START_REF] Manzi | A new facies model for the Upper Gypsum of Sicily (Italy): chronological and paleoenvironmental constraits for the Messinian salinity crisis in the Mediterranean[END_REF], and an approximate duration of 175 kyr for Stage 3 as whole. 2. An alternative tuning by [START_REF] Manzi | A new facies model for the Upper Gypsum of Sicily (Italy): chronological and paleoenvironmental constraits for the Messinian salinity crisis in the Mediterranean[END_REF]Fig. 3a

APPENDIX D

der Laan et al. (2006) on the sedimentary inexpressiveness of the (barely visible) insolation minima peak at ~5.38 Ma. However, these authors considered the insolation minima peaks immediately above and below of too low amplitude to promote the conditions required for gypsum precipitation, but also too high not to have some sedimentary expression. They therefore identified sandstone horizons VI' and VI" as the sedimentary response to these weak insolation/precession signals. The addition of two precessional cycles (i.e. a total of 9) resulted in an astronomical solution that was adjusted one precessional cycle lower than that of Van der Laan et al. (2006), translating into an age of 5.53 Ma for the base of the UG and a total duration of ~200 kyr for Stage 3. But the more conspicuous difference between the two astronomical solutions discussed lies in the timing at which gypsum precipitation occurred, restricted to the 100 kyr eccentricity maxima according to Van der Laan et al. (2006), extended to the 100 kyr eccentricity minima by [START_REF] Manzi | A new facies model for the Upper Gypsum of Sicily (Italy): chronological and paleoenvironmental constraits for the Messinian salinity crisis in the Mediterranean[END_REF].

An age of 5.53 Ma for the first gypsum bed was also obtained by the astronomical tuning of the Upper Gypsum in Cyprus (Manzi et al., 2016a), but there the tuning is performed just by following the recognition, from the base up, of 6 gypsum beds just like in Sicily and therefore arguing for a bed-to-bed correlation with the Sicilian gypsums I-VI. In the consensus model of Roveri et al. (2014a) the base of Stage 3 coincides with the base of the Sicilian UG, placed by [START_REF] Manzi | A new facies model for the Upper Gypsum of Sicily (Italy): chronological and paleoenvironmental constraits for the Messinian salinity crisis in the Mediterranean[END_REF] at 5.55 Ma (Fig. 1A). However, in the model of [START_REF] Manzi | A new facies model for the Upper Gypsum of Sicily (Italy): chronological and paleoenvironmental constraits for the Messinian salinity crisis in the Mediterranean[END_REF] this age is attributed to a cumulate gypsum horizon interpreted as laterally equivalent of the Halite (i.e. Stage 2), and therefore implying the kickoff of Stage 3 at 5.53 Ma (Fig. 3a).

The post-evaporitic successions of the Romagna (Cusercoli and Sapigno sections; [START_REF] Roveri | Stratigraphy of the Messinian post-evaporitic stage in eastern-Romagna (northern Apennines, Italy)[END_REF] and Marche (e.g. Trave and Maccarone sections; [START_REF] Iaccarino | The Trave section (Monte dei Corvi, Ancona, central Italy): an integrated paleontological study of the Messinian deposits[END_REF][START_REF] Cosentino | Refining the Mediterranean "Messinian gap" with high-precision U-Pb zircon geochronology, central and northern Italy[END_REF] areas provided evidence that led to the splitting of Stage 3 into substage 3.1 and 3.2. In the resulting composite section (Roveri et al., 2008a), a shift in the sedimentary facies and stacking pattern is observed (see description in subsection 3.7). Correlation of the sedimentary cyclicity in Romagna was from the (conformable) base of the Pliocene downwards (or from an U-Pb-dated ash layer upward; [START_REF] Cosentino | Refining the Mediterranean "Messinian gap" with high-precision U-Pb zircon geochronology, central and northern Italy[END_REF] and linked three fluvial conglomerates and two black mudstone layers of unknown sedimentological significance to the arid phases of the precession cycles (Fig. 3a; Roveri et al., 2008a). The greater thickness of the oldest conglomerate was possibly assumed to be evocative of an oscillation of the amplitude of the corresponding precession minima peak rather than the amplitude of the peaks responsible for the formation of the other facies. This approach resulted in an age of 5.42 Ma for the first conglomerate (i.e. the substage 3.1/3.2 transition; Fig. 5g) and an approximate duration of 90 kyr for substage 3.2 (the Lago-Mare phase). The same astronomical age is obtained by tuning the Upper Member of the Feos Formation in the Nijar Basin [START_REF] Omodeo-Salé | Tectonic and climatic control on the Late Messinian sedimentary evolution of the Nijar Basin (Betic Cordillera, Southern Spain)[END_REF], where four pelite-conglomerate cycles plus one sandstone capped by the Miocene/Pliocene boundary mark the interval attributed to Stage 3.2 [START_REF] Fortuin | The Messinian of the Nijar basin (SE Spain): sedimentation, depositional environments and paleogeographic evolution[END_REF].

Although the substage 3.1/3.2 transition is linked to a major Mediterranean-scale hydrological re-organization possibly coinciding with the migration of the Paratethyan biota (Roveri et al., 2008a;Grossi et al., 2011), the facies change used for its definition is hardly recognizable elsewhere (see Chapter 3). As such, other tools have been used to equip fragmentary and/or lithological cyclicity-lacking sections with an age model: the (highly controversial) ostracod biozonation (see subsection 5.1; e.g. Stoica et al., 2016;Karakitsios et al., 2017a;Cosentino et al., 2018;[START_REF] Caruso | The late Messinian "Lago-Mare" event and the Zanclean Reflooding in the Mediterranean Sea: new insights from the Cuevas del Almanzora section (Vera Basin, South-Eastern Spain)[END_REF] and the astronomical tuning of magnetic susceptibility records (e.g. Fonte dei Pulcini section, Central Apennines; [START_REF] Cosentino | Orbitally forced paleoenvironmental and paleoclimate changes in the late postevaporitic Messinian of the central Mediterranean Basin[END_REF].

Comparison of Atlantic oxygen isotope records [START_REF] Van Der Laan | Regional climate and glacial control on high-resolution oxygen isotope records from Ain El Beida (latest Miocene, NW Morocco): a cyclostratigraphic analysis in the depth and time domain[END_REF], 2006) and the chronostratigraphy of Roveri et al. (2014a) revealed that Stage 3 sedimentation started during a prominent global eustatic lowstand associated with oxygen isotope (glacial) stage TG12, followed by a latest Messinian deglacial interval which comprised multiple obliquity-and possibly precession-forced global eustatic phases. As documented by [START_REF] Hodell | Correlation of late Miocene to early Pliocene sequences between the Mediterranean and North Atlantic[END_REF] (later revised by [START_REF] Drury | Reinforcing the North Atlantic backbone: revision and extension of the composite splice at ODP Site 982[END_REF], Van der Laan et al. (2006) and Roveri et al. (2014a), the marine replenishment of the Mediterranean did not coincide with any major deglaciation, so non-eustatic causes of the Zanclean megaflood hypothesis are required.

Onshore domain: key sections, sedimentary expression and faunal content

The Alborán region

The westernmost outcrops of Stage 3 deposits in the Mediterranean are located in the Alborán region, close to the present-day Strait of Gibraltar (Fig. 2b). MSC deposits on the margins of this region are poorly developed, possibly because of a late Tortonian uplift that raised the margins above the Mediterranean water level (López-Garrido and Sanz de Galdeano, 1999). Near Malaga, however, two facies associations consisting of m-thick conglomerate-sandstone beds alternating with laminated pelites are documented in the Rio Mendelín section (informally referred to as "LM unit"; [START_REF] Guerra-Merchán | Messinian Lago-Mare deposits near the strait of Gibraltar (Malaga basin, S Spain)[END_REF] and attributed to (part of) the Lago-Mare phase (Fig. 4a) based on their paleontological content. These sediments are squeezed between the Paleozoic basement units, with an erosive contact and associated angular unconformity, and the Pliocene, from which they are separated by another erosional surface draped by conglomeratic accumulations (Fig. 6a). A well-preserved and diverse in situ Paratethyan-type ostracod and molluskan fauna (i.e. Lymnocardiinae and Dreissenidae) typical of shallow waterbodies (up to 100 m deep; [START_REF] Grossi | Late Messinian Lago-Mare ostracods and paleoenvironments of the central and eastern Mediterranean Basin[END_REF][START_REF] Gliozzi | Late Messinian lago-mare ostracod paleoecology: a correspondence analysis aroach[END_REF] with low salinities (5-18‰) is reported from the pelitic units [START_REF] Guerra-Merchán | Messinian Lago-Mare deposits near the strait of Gibraltar (Malaga basin, S Spain)[END_REF]. The overlying Pliocene in the deeper depocenters starts with 30 m-thick littoral conglomerates with marine mollusks passing progressively upwards into deeper water facies, while fan deltas developed at the basin margins (López-Garrido and Sanz de [START_REF] López-Garrido | Neogene sedimentation and tectoniceustatic control of the Malaga basin, South Spain[END_REF][START_REF] Guerra-Merchán | Messinian Lago-Mare deposits near the strait of Gibraltar (Malaga basin, S Spain)[END_REF][START_REF] Abdel-Fattah | Petrophysical characteristics of the messinian abu madi formation in the baltim east and north fields, offshore Nile delta, Egypt[END_REF]. Notably, the overall thickness of the Pliocene deposits reaches 600 m. The detailed regional studies by López-Garrido and Sanz de Galdeano (1999) and [START_REF] Guerra-Merchán | Tectono-sedimentary evolution of the peripheral basins of the Alboran Sea in the arc of Gibraltar during the latest Messinian-Pliocene[END_REF] concluded that accommodation space was created during (Zanclean) sedimentation by local fault-driven subsidence, and that movement on these faults only reversed at the end of the Zanclean causing uplift.

An alternative scenario, based on the finding of (a few) specimens of the nannofossil Ceratolithus acutus, ascribed the LM unit of [START_REF] Guerra-Merchán | Messinian Lago-Mare deposits near the strait of Gibraltar (Malaga basin, S Spain)[END_REF] to the earliest Zanclean (Fig. 4b; [START_REF] Couto | Lago Mare and the Messinian salinity crisis: evidences from the Alboran Sea (S. Spain)[END_REF].

On the southern Alborán margin in Morocco, latest Messinian deposits are reported from the Boudinar and Melilla basins (Fig. 2b). Up to 100 m-thick chaotic deposits containing selenite gypsum fragments, azoic conglomerates, sandstones yielding planktic foraminifera and nannofossils and lacustrine limestones are capped by early Pliocene marine marls [START_REF] Rouchy | Late Messinian to Early Pliocene paleoenvironmental changes in the Melilla Basin (NE Morocco) and their relation to Mediterranean evolution[END_REF][START_REF] Azdimousa | Geodynamique des bordures meridionales de la mer d'Alboran; alication de la stratigraphie sequentielle dans le bassin neogene de Boudinar Rif oriental[END_REF][START_REF] Cornée | The Messinian erosional surface and early Pliocene reflooding in the Alboran Sea: new insights from the Boudinar basin[END_REF][START_REF] Merzeraud | Sedimentology and sequence stratigraphy of the late-Messinian-Early pliocene continental to marine deposits of the Boudinar basin (North Morocco)[END_REF]. Due to their stratigraphic position, these continental to lacustrine deposits are interpreted as the local expression of the Lago-Mare phase [START_REF] Cornée | The Messinian erosional surface and early Pliocene reflooding in the Alboran Sea: new insights from the Boudinar basin[END_REF] or alternatively as Zanclean successions [START_REF] Azdimousa | Geodynamique des bordures meridionales de la mer d'Alboran; alication de la stratigraphie sequentielle dans le bassin neogene de Boudinar Rif oriental[END_REF].

Algeria

The Chelif Basin in Algeria (Fig. 2b) displays the typical marginal Messinian succession comprising Tortonian to lower Messinian blue marls, diatomite-bearing sediments (Tripoli unit), cyclically-arranged primary evaporites (13 couplets), ostracod-rich post-evaporitic deposits and Zanclean foraminiferal marls [START_REF] Rouchy | The end of the Messinian salinity crisis: evidences from the Chelif Basin (Algeria)[END_REF]. The postevaporitic sediments show a great lateral variability in both thickness (from few meters up to 125 m) and facies. They are mainly dominated by F. Andreetto et al. terrigenous clastic lithologies, associated in the marginal areas with sandy carbonates and stromatolitic limestones. A mixed faunal assemblage of non-marine (Paratethyan-like ostracods) and marine (benthic and planktic foraminifera) organisms is present, showing an increase in ostracod species diversity from the bottom to the top [START_REF] Rouchy | The end of the Messinian salinity crisis: evidences from the Chelif Basin (Algeria)[END_REF].

Neogene basins of the Eastern Betics (Spain)

The external Neogene basins (Sorbas, Nijar, Vera and Bajo Segura) of the eastern Betic Cordillera (SE Spain; Fig. 2b) represent an important laboratory for understanding Messinian events. In particular, the Sorbas and Nijar basins preserve two allegedly continuous successions spanning the entire MSC (e.g. [START_REF] Roveri | The Terminal Carbonate Complex: the record of sea-level changes during the Messinian salinity crisis[END_REF][START_REF] Omodeo-Salé | Tectonic and climatic control on the Late Messinian sedimentary evolution of the Nijar Basin (Betic Cordillera, Southern Spain)[END_REF]. The two basins are similar in many respects. Their stratigraphic organization, for example, suggests they were connected for much of the late Miocene up until MSC Stage 1 [START_REF] Fortuin | The Messinian of the Nijar basin (SE Spain): sedimentation, depositional environments and paleogeographic evolution[END_REF], which is represented by the gypsiferous Yesares Member (e.g. Lu, 2006). However, facies differences are prominent in the Stage 3 formations according to the chronostratigraphic frameworks of [START_REF] Roveri | The Terminal Carbonate Complex: the record of sea-level changes during the Messinian salinity crisis[END_REF] for the Sorbas Basin and Omodeo-Salé et al. ( 2012) for the Nijar Basin (Fig. 4a). Lithostratigraphically, two members are discerned between the Yesares Member and the basal Zanclean: the Sorbas and Zorreras members in the Sorbas Basin (Figs. 4a,5a) and the lower and upper members of the Feos Fm. in Nijar (Figs. 3a, 4a;[START_REF] Roep | Late highstand patterns of shifting and steing coastal barriers and washover-fans (Late Messinian, Sorbas Basina, SE Spain)[END_REF][START_REF] Krijgsman | Astrochronology for the Messinian Sorbas basin (SE Spain) and orbital (precessional) forcing for evaporite cyclicity[END_REF][START_REF] Fortuin | The Messinian of the Nijar basin (SE Spain): sedimentation, depositional environments and paleogeographic evolution[END_REF][START_REF] Braga | Testing models for the Messinian salinity crisis: the Messinian record in Almería, SE Spain[END_REF][START_REF] Roveri | The Terminal Carbonate Complex: the record of sea-level changes during the Messinian salinity crisis[END_REF]Roveri et al., , 2019a;;[START_REF] Omodeo-Salé | Tectonic and climatic control on the Late Messinian sedimentary evolution of the Nijar Basin (Betic Cordillera, Southern Spain)[END_REF].

The Sorbas Member (see [START_REF] Roep | Late highstand patterns of shifting and steing coastal barriers and washover-fans (Late Messinian, Sorbas Basina, SE Spain)[END_REF][START_REF] Aufgebauer | Messinian to Pliocene transition in the deep part of the Sorbas Basin, SE Spain-a new description of the depositional environment during the Messinian Salinity Crisis[END_REF][START_REF] Aufgebauer | Messinian to Pliocene transition in the deep part of the Sorbas Basin, SE Spain-a new description of the depositional environment during the Messinian Salinity Crisis[END_REF] for a more detailed sedimentological description) consists of three overlapping coarsening-upward depositional sequences made of offshore clays and marls passing upward into shelf muds and coastal sandstone bodies. Still unclear is the chemistry of the subaqueous environment during the formation of the Sorbas Member and the provenance of the water fluxes. These shallow-water deposits are conformably replaced upward by the Zorreras Member that comprises alternations of reddish siltstones and sandstones (Fig. 5a) organized in five (or eight) lithological cycles expressing continental environments [START_REF] Martín-Suárez | On the age of the continental deposits of the Zorreras Member (Sorbas Basin, SE Spain)[END_REF][START_REF] Aufgebauer | Messinian to Pliocene transition in the deep part of the Sorbas Basin, SE Spain-a new description of the depositional environment during the Messinian Salinity Crisis[END_REF]. Up to four lenticular white limestone beds bearing brackish Paratethyan-like ostracods (Cyprideis, Loxocorniculina djafarovi and freshwater species of the family Limnocytheridae), bivalves and Chara oogonia [START_REF] Andreetto | Sedimentological and ostracodological observations on Messinian post-evaporite deposits in some southeastern Spanish basins[END_REF][START_REF] Aufgebauer | Messinian to Pliocene transition in the deep part of the Sorbas Basin, SE Spain-a new description of the depositional environment during the Messinian Salinity Crisis[END_REF] are found interrupting the fluviatile sequence (Fig. 5a) and are linked to either episodic flooding by local rivers [START_REF] Braga | Testing models for the Messinian salinity crisis: the Messinian record in Almería, SE Spain[END_REF][START_REF] Aufgebauer | Messinian to Pliocene transition in the deep part of the Sorbas Basin, SE Spain-a new description of the depositional environment during the Messinian Salinity Crisis[END_REF] or episodic Mediterranean ingressions [START_REF] Fortuin | The Messinian of the Nijar basin (SE Spain): sedimentation, depositional environments and paleogeographic evolution[END_REF]Andreetto et al., 2021). A correct interpretation of the paleodepositional environment of these limestone beds is crucial for the discussion concerning the Mediterranean base-level position during the Lago-Mare phase. In fact, if the Sorbas Basin was relatively shallow during Zorreras deposition (50-100 m; Roveri et al., 2019a[START_REF] Roveri | Stratigraphic relationships between shallow-water carbonates and primary gypsum: insights from the Messinian succession of the Sorbas Basin (Betic Cordillera, Southern Spain)[END_REF], repeated and sudden Mediterranean incursions would indicate that the Mediterranean Basin was relatively full and that its base level was oscillating, possibly with precessional periodicity (Andreetto et al., 2021). The contact between the Zorreras Mb. and the overlying nearshore Pliocene (<50 m depositional paleodepth; Roveri et al., 2019a) in the Sorbas Basin is conformable and expressed differently around the basin, ranging from a bivalves-rich bed overlain by a yellow, fossiliferous calcarenite floored by a gravelly lag deposit [START_REF] Mather | A Field Guide to the Neogene Sedimentary Basins of the Almería Province[END_REF] to a grey marl horizon with marine foraminifera assemblages followed by a second shell-rich bed (Roveri et al., 2019a). Similar to the situation in Malaga, the rare identification of Ceratolithus acutus in sediments of the continental Zorreras Mb. led [START_REF] Clauzon | New insights on the Sorbas Basin (SE Spain): the onshore reference of the Messinian salinity crisis[END_REF] to put forward an alternative chronostratigraphic and paleoenvironmental interpretation for the Sorbas MSC succession, shifting the Zorreras Mb. into the Pliocene (Fig. 4b) and thus associating the presence of brackish Paratethyanlike ostracods with exchanges between the Mediterranean and Paratethys following the Mediterranean re-filling, at high sea level.

In the Nijar Basin (Fig. 2b), the latest Messinian Feos Formation is bracketed at the base and top by an erosional surface along the basin margins and its correlative conformity in the deeper parts (Fig. 3a; [START_REF] Fortuin | The Messinian of the Nijar basin (SE Spain): sedimentation, depositional environments and paleogeographic evolution[END_REF][START_REF] Aguirre | The Messinian post-evaporitic deposits of the Gafares area (Almeria-Nijar basin, SE Spain). A new view of the "Lago-Mare" facies[END_REF][START_REF] Omodeo-Salé | Tectonic and climatic control on the Late Messinian sedimentary evolution of the Nijar Basin (Betic Cordillera, Southern Spain)[END_REF]. The Lower Feos Member consists of azoic, graded and locally slumped siliciclastic-carbonate beds alternating with gypsarenites and gypsiltites and including a laterally continuous Mnrich bed [START_REF] Fortuin | The Messinian of the Nijar basin (SE Spain): sedimentation, depositional environments and paleogeographic evolution[END_REF][START_REF] Omodeo-Salé | Tectonic and climatic control on the Late Messinian sedimentary evolution of the Nijar Basin (Betic Cordillera, Southern Spain)[END_REF]. In the basin center (e.g. Barranco de los Castellones section; Fig. 3a) the Upper Feos member comprises four complete lithological cycles of mthick conglomerate to sandstone beds alternating with laminated pelites (Fig. 5b), and one incomplete cycle, which only consists of a sandstone horizon conformably capped by the Pliocene Cuevas Fm. (Fig. 6b; [START_REF] Andreetto | Alternate interpretation of the Messinian salinity crisis, controversy resolved?[END_REF][START_REF] Clauzon | Influence of Mediterranean sea-level changes on the Dacic Basin (Eastern Paratethys) during the late Neogene: the Mediterranean Lago Mare facies deciphered[END_REF], which envisaged two Lago-Mare episodes (LM1 and LM3) that occurred before and after the main Mediterranean drawdown event, during which LM2 was deposited in the deep desiccated basins [START_REF] Couto | Lago Mare and the Messinian salinity crisis: evidences from the Alboran Sea (S. Spain)[END_REF][START_REF] Popescu | Lago Mare episodes around the Messinian-Zanclean boundary in the deep southwestern Mediterranean[END_REF]see Roveri et al., 2008c and[START_REF] Andreetto | Migration of the dinoflagellate Galeacysta etrusca and its implications for the Messinian Salinity Crisis[END_REF] for further explanations). Note, in (b), the shifting of the position of the main erosional phase in Sicily through time as well as the time of the marine replenishment in the Apennines. [START_REF] Fortuin | The Messinian of the Nijar basin (SE Spain): sedimentation, depositional environments and paleogeographic evolution[END_REF]. A rich fauna of mixed brackish ostracods and marine foraminifera is found in all four pelitic beds [START_REF] Bassetti | Evolution of depositional environments after the end of Messinian Salinity Crisis in Nijar basin (SE Betic Cordillera)[END_REF]. Its origin is questionable. These ostracods were regarded as endemic to the Mediterranean and inhabiting endorheic lakes by [START_REF] Bassetti | Evolution of depositional environments after the end of Messinian Salinity Crisis in Nijar basin (SE Betic Cordillera)[END_REF]. However, later they were shown to have been misidentified and were instead considered Paratethys-derived by Stoica et al. (2016;see subsection 5.1). Planktonic and deep-water benthic foraminifera are widely considered reworked by [START_REF] Fortuin | The Messinian of the Nijar basin (SE Spain): sedimentation, depositional environments and paleogeographic evolution[END_REF], [START_REF] Bassetti | Evolution of depositional environments after the end of Messinian Salinity Crisis in Nijar basin (SE Betic Cordillera)[END_REF] and [START_REF] Omodeo-Salé | Tectonic and climatic control on the Late Messinian sedimentary evolution of the Nijar Basin (Betic Cordillera, Southern Spain)[END_REF], in place by [START_REF] Aguirre | The Messinian post-evaporitic deposits of the Gafares area (Almeria-Nijar basin, SE Spain). A new view of the "Lago-Mare" facies[END_REF].

In the Vera Basin (Fig. 2b), in situ gypsum deposits are missing because of widespread erosion or non-deposition and MSC deposits are only represented by ~12 m of laminated varicolored marly clays (Unit 2 Fig. 4a), which are best exposed in the Cuevas del Almanzora section [START_REF] Fortuin | The enigmatic Messinian -Pliocene section of Cuevas del Almanzora (Vera Basin, SE Spain) revisited-erosional features and strontium isotope ages[END_REF]Fig. 5c). These clays contain a well-preserved and diversified in situ fauna of Paratethyan-like ostracod and shallow-water, benthic foraminifera mixed with physically reworked (mostly from the lower Messinian Abad marls) planktic and deep-water benthic foraminifera [START_REF] Fortuin | The enigmatic Messinian -Pliocene section of Cuevas del Almanzora (Vera Basin, SE Spain) revisited-erosional features and strontium isotope ages[END_REF]Stoica et al., 2016;[START_REF] Caruso | The late Messinian "Lago-Mare" event and the Zanclean Reflooding in the Mediterranean Sea: new insights from the Cuevas del Almanzora section (Vera Basin, South-Eastern Spain)[END_REF]. The marly clays are assigned by Stoica et al. (2016) and [START_REF] Caruso | The late Messinian "Lago-Mare" event and the Zanclean Reflooding in the Mediterranean Sea: new insights from the Cuevas del Almanzora section (Vera Basin, South-Eastern Spain)[END_REF] to (roughly) the whole late Messinian Lago-Mare phase (Fig. 4a) based on the ostracod biozonation of Grossi et al. (2011) and are considered to represent either sedimentation in an isolated lake subject to base-level and salinity fluctuations [START_REF] Caruso | The late Messinian "Lago-Mare" event and the Zanclean Reflooding in the Mediterranean Sea: new insights from the Cuevas del Almanzora section (Vera Basin, South-Eastern Spain)[END_REF] or deposition in a coastal lagoon that was connected to the water mass filling the open Mediterranean (Stoica et al., 2016;Andreetto et al., 2021). Similar to Malaga, these sediments are topped by an erosive surface draped by a conglomeratic accumulation which is overlain by the open marine fauna-rich sediments of the basal Zanclean [START_REF] Fortuin | The enigmatic Messinian -Pliocene section of Cuevas del Almanzora (Vera Basin, SE Spain) revisited-erosional features and strontium isotope ages[END_REF][START_REF] Caruso | The late Messinian "Lago-Mare" event and the Zanclean Reflooding in the Mediterranean Sea: new insights from the Cuevas del Almanzora section (Vera Basin, South-Eastern Spain)[END_REF]. This erosion feature likely indicates that the Miocene/ Pliocene transition followed a base-level lowstand in the Vera Basin.

Stage 3 deposits (Garrucha Fm.) in the easternmost basin of the Betic Cordillera, the Bajo Segura Basin (Fig. 2b), are bounded below and above by two erosional surfaces related to lowered Mediterranean base-levels and discontinuously present due to the widespread fluvial erosion that occurred at the Miocene/Pliocene boundary [START_REF] Soria | The stratigraphic record of the Messinian salinity crisis in the northern margin of the Bajo Segura basin (SE Spain)[END_REF](Soria et al., , 2008a(Soria et al., , 2008b)). The Garrucha Fm. shows a maximum thickness of 100 m in its type section [START_REF] Soria | Estratigrafía y biomagnetoestratigrafía del Messiniense en la sección del Garruchal (Cuenca del Bajo Segura)[END_REF](Soria et al., , 2008b)). It consists of 20-50 cm thick sandstone bodies interrupting a dominantly marly succession deposited in a subaqueous environment inhabited by Cyprideis sp. and euryhaline, shallow-water benthic foraminifera (Ammonia beccarii, Elphidium granosum, Elphidium macellum, Haynesina germanica and Quinqueloculina laevigata). Planktic foraminifera are also observed and for a long time were considered to be physically reworked [START_REF] Soria | The stratigraphic record of the Messinian salinity crisis in the northern margin of the Bajo Segura basin (SE Spain)[END_REF](Soria et al., , 2008b)). However, some stratigraphic levels contain dwarf tests of long-ranging taxa such as Globoturborotalita decoraperta, Globigerina bulloides, and Neogloboquadrina spp. which recently have been interpreted as being in-situ mostly due to the absence of notable signs of reworking (Corbí and Soria, 2016). Among these dwarf taxa is Neogloboquadrina acostaensis (dextral; Corbí and Soria, 2016). Since this group is mainly dextral in the latest Messinian Atlantic successions (e.g. [START_REF] Sierro | Late Miocene globorotaliid event-stratigraphy and biogeography in the NE-Atlantic and Mediterranean[END_REF][START_REF] Bassetti | Evolution of depositional environments after the end of Messinian Salinity Crisis in Nijar basin (SE Betic Cordillera)[END_REF], this may indicate that Atlantic inflow to the Mediterranean occurred during the late Messinian and the base level of the Mediterranean was high enough to reach the marginal Bajo Segura Basin. The Miocene/Pliocene boundary is, once again, marked by an erosional surface which outlines up to 200 m deep paleovalleys engraved down into the pre-MSC sediments and filled with conglomerates and sandstones of claimed coastal and shallow marine environments [START_REF] Soria | The stratigraphic record of the Messinian salinity crisis in the northern margin of the Bajo Segura basin (SE Spain)[END_REF](Soria et al., , 2008b;;[START_REF] García-García | Architecture analysis of a river flood-dominated delta during an overall sea-level rise (Early Pliocene, SE Spain)[END_REF]Corbí et al., 2016).

Mallorca

Mallorca, which constitutes an emerged segment of the Balearic Promontory (Fig. 2), does not expose the classical MSC evaporite sequence. Instead, two main MSC-related units are found above late Tortonian-Messinian reefal carbonates (Reef Complex Unit) and beneath the Pliocene: the Santanyí limestones and the Ses Olles Formation (Mas and Fornós, 2020 and references therein). The Santanyí limestones are microbialites and oolite-dominated sediments in which a baleen whale neurocranium has been found (Mas et al., 2018a). This unit was interpreted either as a Terminal Carbonate Complex (TCC) laterally equivalent to the Primary Lower Gypsum (PLG) which has been drilled in the deeper parts of the bay of Palma [START_REF] Mas | The messinian salinity crisis in Mallorca: New insights for a western mediterranean stratigraphic scenario[END_REF] or as timeequivalent to the Reef Complex Unit (e.g. [START_REF] Arenas | Microbial deposits in upper Miocene carbonates, Mallorca, Spain[END_REF][START_REF] Suárez-González | Interplay between biotic and environmental conditions in presalt Messinian microbialites of the western Mediterranean (Upper Miocene, Mallorca, Spain)[END_REF]. The Ses Olles Formation consists of marls, sandy-marls and marly-calcareous lacustrine deposits rich in in-situ freshwater Chara spp., brackish water Paratethyan-like mollusks and ostracods and littoral benthic foraminifera (Elphidium sp., Ammonia sp.).

The upper contact of the Ses Olles Formation with the Pliocene corresponds to an erosional ravinement surface draped by a transgressive lag of coastal deposits usually containing coquinas and/or conglomerates [START_REF] Mas | Definició i caracterització de la Formació ses Olles (Lago mare, messinià terminal) a l'Illa de Mallorca (illes balears, mediterrània Occidental)[END_REF][START_REF] Mas | El registre estratigr`afic del Messini`a terminal i del Pliocè a l'illa de Mallorca[END_REF][START_REF] Mas | The messinian salinity crisis in Mallorca: New insights for a western mediterranean stratigraphic scenario[END_REF]. The lower contact of the Ses Olles Formation with the Santanyí limestones is sporadically marked by a well-developed reddish paleosol [START_REF] Mas | Definició i caracterització de la Formació ses Olles (Lago mare, messinià terminal) a l'Illa de Mallorca (illes balears, mediterrània Occidental)[END_REF][START_REF] Mas | El registre estratigr`afic del Messini`a terminal i del Pliocè a l'illa de Mallorca[END_REF][START_REF] Mas | The messinian salinity crisis in Mallorca: New insights for a western mediterranean stratigraphic scenario[END_REF], which indicates that a (unquantified) period of subaerial exposure occurred before the emplacement of the Ses Olles Fm. However, in their more recent study, [START_REF] Mas | The messinian salinity crisis in Mallorca: New insights for a western mediterranean stratigraphic scenario[END_REF] surprisingly conclude that the Ses Olles Formation has a conformable contact with the Santanyí limestones, ascribed to part of Stage 1. This led [START_REF] Mas | The messinian salinity crisis in Mallorca: New insights for a western mediterranean stratigraphic scenario[END_REF] to conclude that the emplacement of the Ses Olles Fm. pre-dated the MSC peak and that the erosional surface marking the Miocene/Pliocene boundary is associated with a 270 kyr hiatus linked to the main MSC base-level drawdown (Fig. 4b). This conclusion is, however, in disagreement with the unconformity at the base of the Ses Olles Fm., which instead points to the deposition of the Ses Olles Fm. (and therefore to the arrival of the Paratethyan fauna in Mallorca) at some point during Stage 3 of Roveri et al. (2014a).

Piedmont Basin

The Piedmont Basin (NW Italy) contains the northernmost record of the MSC (Fig. 2b). The terminal MSC sediments (i.e. the Cassano Spinola Conglomerates Fm.) overlay pre-MSC units, the PLG deposits (Gessoso Solfifera Fm.) or reworked evaporites (Valle Versa chaotic complex, VVC) and underly the Zanclean marls of the Argille Azzurre Fm. (Dela Pierre et al., 2011).

The Cassano Spinola Conglomerates is splitted in two sub-units by Dela [START_REF] Pierre | The northermost record of the Messinian salinity crisis (Piedmont Basin, NW Italy)[END_REF]. Sub-unit a consists of azoic grey mudstones turning to yellowish silty mudstones (Fig. 5d) typified by in situ root traces, paleosols and mud cracks and including three/four intercalated lens-shaped, cross-bedded conglomeratic layers [START_REF] Ghibaudo | Litostratigrafia, sedimentologia ed evoluzione tettonico-sedimentaria dei depositi miocenici del margine sud-orientale del bacino terziario ligure-piemontese (Valli Borbera, Scrivia e Lemme). In memoria di Carlo Sturani[END_REF]Dela Pierre et al., 2011, 2016). Abundant land plant leaves and a diverse terrestrial vertebrate fauna are found in the yellowish siltstones, which have been interpreted as overbank deposits [START_REF] Harzhauser | Late Messinian continental and Lago-Mare gastropods from the Tertiary Piedmont Basin[END_REF]Colombero et al., 2017 and references therein). In this continental interval, a low-diversity fish fauna consisting of otoliths of marine and Paratethyan species is found [START_REF] Grunert | Estuarine Lago Mare fauna from the Tertiary Piedmont Basin indicates episodic Atlantic/Mediterranean exchange during the final stage of the Mediterranean Salinity Crisis[END_REF][START_REF] Carnevale | Fossil marine fishes and the "Lago Mare" event: Has the Mediterranean ever transformed into a brackish lake?[END_REF][START_REF] Schwarzhans | Late Miocene-Early Pliocene evolution of Mediterranean gobies and their environmental and biogeographic significance[END_REF]. These otoliths were Sr-dated to the early-middle Miocene [START_REF] Grunert | Estuarine Lago Mare fauna from the Tertiary Piedmont Basin indicates episodic Atlantic/Mediterranean exchange during the final stage of the Mediterranean Salinity Crisis[END_REF]. Nevertheless, they were concluded not to be physically reworked, but rather to have been transported by large marine predators, therefore implying a Piedmont Basin-(marine) Mediterranean connection was in force [START_REF] Grunert | Estuarine Lago Mare fauna from the Tertiary Piedmont Basin indicates episodic Atlantic/Mediterranean exchange during the final stage of the Mediterranean Salinity Crisis[END_REF]; see subsection 5.6). Sub-unit b (i.e. Strati a Congeria sensuSturani, 1973) is made of grey mudstones bearing a mixture of in-situ brackish water mollusks [START_REF] Sturani | A fossil eel (Anguilla sp.) from the Messinian of Alba (Tertiary Piedmont Basin). Paleoenvironmental and paleogeographic implications[END_REF][START_REF] Esu | Latest Messinian "Lago-Mare" Lymnocardiinae from Italy: close relations with the Pontian fauna from the Dacic Basin[END_REF] and ostracods [START_REF] Trenkwalder | The Miocene/Pliocene boundary in the Early Pliocene micropaleontological record: new data from the Tertiary Piedmont Basin (Moncucco quarry, Torino Hill, northwestern Italy)[END_REF] of Paratethyan affinity along with physically reworked foraminifera and calcareous nannofossils [START_REF] Trenkwalder | The Miocene/Pliocene boundary in the Early Pliocene micropaleontological record: new data from the Tertiary Piedmont Basin (Moncucco quarry, Torino Hill, northwestern Italy)[END_REF][START_REF] Violanti | Micropaleontological analyses of the Narzole core: biostratigraphy and paleoenvironment of the late Messinian and early Zanclean of Piedmont (Northwestern Italy)[END_REF]. The transition to the Pliocene Argille Azzurre Fm. is sharp above a characteristic black and azoic sandy layer (Fig. 6c) rich in terrigenous and intrabasinal (i.e., glaucony and phosphates) grains and disarticulated valves of both brackish-water and continental bivalves, but barren of in-situ fossils [START_REF] Trenkwalder | The Miocene/Pliocene boundary in the Early Pliocene micropaleontological record: new data from the Tertiary Piedmont Basin (Moncucco quarry, Torino Hill, northwestern Italy)[END_REF]. The occurrence, at its top and directly below the Argille Azzurre Fm., of abundant Thalassinoides trace fossils filled with Pliocene sediments led [START_REF] Trenkwalder | The Miocene/Pliocene boundary in the Early Pliocene micropaleontological record: new data from the Tertiary Piedmont Basin (Moncucco quarry, Torino Hill, northwestern Italy)[END_REF] and Dela [START_REF] Pierre | The northermost record of the Messinian salinity crisis (Piedmont Basin, NW Italy)[END_REF] to interpret the top surface of this layer as an omission surface. This surface indicates a period of basin starvation (and therefore a hiatus) due to a sudden increase in water-depth, ascribed by [START_REF] Trenkwalder | The Miocene/Pliocene boundary in the Early Pliocene micropaleontological record: new data from the Tertiary Piedmont Basin (Moncucco quarry, Torino Hill, northwestern Italy)[END_REF] to the Zanclean reflooding. This hiatus may have lasted for only part of the late Messinian [START_REF] Violanti | Micropaleontological analyses of the Narzole core: biostratigraphy and paleoenvironment of the late Messinian and early Zanclean of Piedmont (Northwestern Italy)[END_REF][START_REF] Pierre | The northermost record of the Messinian salinity crisis (Piedmont Basin, NW Italy)[END_REF] or may have endured into the Pliocene [START_REF] Trenkwalder | The Miocene/Pliocene boundary in the Early Pliocene micropaleontological record: new data from the Tertiary Piedmont Basin (Moncucco quarry, Torino Hill, northwestern Italy)[END_REF].

Po Plain

To the east, the Messinian sediments in the Piedmont Basin disappear beneath the km-thick Plio-Quaternary succession of the Po Plain-Adriatic Foredeep (PPAF; Fig. 2a). By definition of [START_REF] Ghielmi | Sedimentary and tectonic evolution in the eastern Po-Plain and northern Adriatic Sea area from Messinian to Middle Pleistocene (Italy)[END_REF] and [START_REF] Amadori | Restored topography of the Po Plain-Northern Adriatic region during the Messinian base level drop-Implications for the physiography and compartmentalization of the paleo-Mediterranean basin[END_REF], the PPAF includes two main elongated depocenters enclosed within the northern Apennines to the South and the Southern Alps to the North: the easternmost portion of the Po Plain and the whole present-day northern Adriatic Sea. Here, for simplicity, we include in the definition of PPAF also its westernmost depocenters of the Western Po Plain Foredeep.

The Messinian-Pleistocene sedimentary sequence, studied through the integration of seismic and borehole observations, is mostly represented by thick sequences of turbidite deposits in the foreland depocenter passing, towards the margins, to fluvial and deltaic systems related to the proximity of the marginal thrust-fold-belts [START_REF] Cipollari | Extension-and compression-related basins in central Italy during the Messinian Lago-Mare event[END_REF][START_REF] Ghielmi | Sedimentary and tectonic evolution in the eastern Po-Plain and northern Adriatic Sea area from Messinian to Middle Pleistocene (Italy)[END_REF][START_REF] Ghielmi | Late Miocene-Middle Pleistocene sequences in the Po Plain-Northern Adriatic Sea (Italy): the stratigraphic record of modification phases affecting a complex foreland basin[END_REF]Rossi et al., 2015a;[START_REF] Rossi | Outcrop and seismic expression of stratigraphic patterns driven by accommodation and sediment suly turnarounds: Implications on the meaning and variability of unconformities in syn-orogenic basins[END_REF]. During MSC Stage 1, primary evaporites and dolomicrites were deposited in some shallow-water settings, while evaporitic deposition was inhibited in the deep-water settings, where it was replaced by deposition of anoxic mudstones [START_REF] Ghielmi | Sedimentary and tectonic evolution in the eastern Po-Plain and northern Adriatic Sea area from Messinian to Middle Pleistocene (Italy)[END_REF]. Instead, the post-evaporitic deposits consist of large thicknesses (up to 1 km) and volumes of coarsegrained clastics (LM1 and LM2 of [START_REF] Rossi | Relative sea-level changes, local tectonic setting and basin margin sedimentation in the interference zone between two orogenic belts: seismic stratigraphic examples from Padan foreland basin, northern Italy[END_REF]ME3 or Fusignano Fm. of Ghielmi et al., 2010;ME4 of Ghielmi et al., 2013; ME3b and possibly ME3a of Rossi et al., 2015a). Several authors [START_REF] Ghielmi | Sedimentary and tectonic evolution in the eastern Po-Plain and northern Adriatic Sea area from Messinian to Middle Pleistocene (Italy)[END_REF][START_REF] Ghielmi | Late Miocene-Middle Pleistocene sequences in the Po Plain-Northern Adriatic Sea (Italy): the stratigraphic record of modification phases affecting a complex foreland basin[END_REF]Rossi et al., 2015a;[START_REF] Amadori | Restored topography of the Po Plain-Northern Adriatic region during the Messinian base level drop-Implications for the physiography and compartmentalization of the paleo-Mediterranean basin[END_REF][START_REF] Cazzini | New seismic evidence of the Messinian paleomorphology beneath Lake Maggiore area (Italy)[END_REF] showed that these post-evaporitic sediments are the infilling of ca. N-S and NW-SE trending, V-shaped valleys (Fig. 5e). These valleys were carved at least as far as 50 km into the Alps, to a depth up to 1 km into the pre-and syn-evaporitic Messinian deposits and nicely shape the present-day river network of the southern Alps [START_REF] Amadori | Restored topography of the Po Plain-Northern Adriatic region during the Messinian base level drop-Implications for the physiography and compartmentalization of the paleo-Mediterranean basin[END_REF].

Different mechanisms for the incision have been proposed, with major implications for the desiccated vs full Mediterranean controversy (Figs. 1b-c). [START_REF] Ghielmi | Sedimentary and tectonic evolution in the eastern Po-Plain and northern Adriatic Sea area from Messinian to Middle Pleistocene (Italy)[END_REF][START_REF] Ghielmi | Late Miocene-Middle Pleistocene sequences in the Po Plain-Northern Adriatic Sea (Italy): the stratigraphic record of modification phases affecting a complex foreland basin[END_REF], Rossi et al. (2015a), [START_REF] Amadori | Restored topography of the Po Plain-Northern Adriatic region during the Messinian base level drop-Implications for the physiography and compartmentalization of the paleo-Mediterranean basin[END_REF] and [START_REF] Cazzini | New seismic evidence of the Messinian paleomorphology beneath Lake Maggiore area (Italy)[END_REF] ascribed the valley incision along the PPAF northern margin to fluvial erosion, whose basinward shifting was triggered by the Stage 2 Mediterranean drawdown, estimated to have been around 800-900 m [START_REF] Ghielmi | Late Miocene-Middle Pleistocene sequences in the Po Plain-Northern Adriatic Sea (Italy): the stratigraphic record of modification phases affecting a complex foreland basin[END_REF][START_REF] Amadori | Restored topography of the Po Plain-Northern Adriatic region during the Messinian base level drop-Implications for the physiography and compartmentalization of the paleo-Mediterranean basin[END_REF]. In this case, Stage 3 deposition in the PPAF occurred in endorheic lakes fed by the Alpine rivers and kept isolated until the Zanclean, when the sudden sea-level rise following the Zanclean reflooding was enough to bypass morphological highs (e.g. Gargano-Pelagosa and/or Otranto paleosills) located in the southern Adriatic foredeep (Fig. 2a; see Pellen et al., 2017;[START_REF] Amadori | Restored topography of the Po Plain-Northern Adriatic region during the Messinian base level drop-Implications for the physiography and compartmentalization of the paleo-Mediterranean basin[END_REF]Manzi et al., 2020). Conversely, [START_REF] Winterberg | Messinian or Pleistocene valley incision within the Southern Alps[END_REF] suggested that the over-deepened valleys on the southern slope of the Alps are related to Pleistocene glacial erosion. Although [START_REF] Winterberg | Messinian or Pleistocene valley incision within the Southern Alps[END_REF] do not address the paleoenvironment during the Messinian, this interpretation does not rule out the possibility that (at least part of) Stage 3 sedimentation occurred in a PPAF connected to the Mediterranean water mass and that no catastrophic reflooding occurred at the Miocene/Pliocene boundary. The conclusion of a non-catastrophic refilling was also drawn by Pellen et al. (2017) on the basis of the onshore Adriatic record (see subsection 3.7).

F. Andreetto et al.

Apennine system

The Messinian deposits resurface to the south of the PPAF sector and extensive sections are found in several basins on both the foreland domain (Adriatic side of the partially uplifted Apennine chain), subjected to compressional tectonics during the late Messinian, and the back-arc domain (Tyrrhenian side), contemporaneously affected by extension (Fig. 2b; [START_REF] Cipollari | Extension-and compression-related basins in central Italy during the Messinian Lago-Mare event[END_REF][START_REF] Schildgen | Sea level and climate forcing of the Sr isotope composition of Late Miocene Mediterranean marine basins[END_REF]Cosentino et al., 2018). Overall, the MSC record of the Apennines is subdivided into an evaporitic and post-evaporitic interval squeezed in between two marine units (Messinian Euxinic Shales Fm. at the base and Zanclean Argille Azzurre Fm. atop; Fig. 4a). Different vertical motions related to ongoing Apenninic tectonics resulted in the deposition of Stage 3 sediments with highly variable sedimentary expression and stratigraphic resolution from basin to basin. The post-evaporitic deposits are alternatively found resting unconformably, with an erosional contact associated to an angular unconformity, above the alternations of the Gessoso Solfifera Fm./PLG, or conformably above evaporitic-free cycles lateral equivalent of the marginal PLG (Fig. 4a; e.g. [START_REF] Roveri | Stratigraphy of the Messinian post-evaporitic stage in eastern-Romagna (northern Apennines, Italy)[END_REF]Roveri et al., , 2008a)). This led to the conclusion that both shallow and deep-water successions are present in the Apennine foredeep system [START_REF] Roveri | The Mediterranean Messinian salinity crisis: an Apennines foredeep perspective[END_REF].

The physical-stratigraphic model developed for the post-evaporitic interval in the Romagna area (i.e. Northern Apennines) and applied to the whole Apennine domain was subdivided into two allounits (named p-ev 1 and p-ev 2 ) based on a basin-wide shift in facies, overall stacking patterns and depositional trends (i.e. progradational and retrogradational, respectively; [START_REF] Roveri | Stratigraphy of the Messinian post-evaporitic stage in eastern-Romagna (northern Apennines, Italy)[END_REF][START_REF] Roveri | The Mediterranean Messinian salinity crisis: an Apennines foredeep perspective[END_REF][START_REF] Roveri | The Adriatic foreland record of Messinian events (central Adriatic sea, Italy)[END_REF]Roveri et al., , 2008a;;[START_REF] Manzi | Deep-water clastic evaporites deposition in the Messinian Adriatic foredeep (northern Apennines, Italy): did the Mediterranean ever dry out?[END_REF][START_REF] Manzi | The deep-water counterpart of the messinian lower evaporites in the apennine foredeep: the fanantello section (northern Apennines, Italy)[END_REF]Manzi et al., , 2020)). Allounit p-ev 1 only accumulated in deep-water settings (e.g. Cusercoli, Sapigno, Maccarone and Trave sections; [START_REF] Roveri | Stratigraphy of the Messinian post-evaporitic stage in eastern-Romagna (northern Apennines, Italy)[END_REF][START_REF] Iaccarino | The Trave section (Monte dei Corvi, Ancona, central Italy): an integrated paleontological study of the Messinian deposits[END_REF][START_REF] Cosentino | Refining the Mediterranean "Messinian gap" with high-precision U-Pb zircon geochronology, central and northern Italy[END_REF] during the subaerial exposure of the basin margins (e.g. Vena del Gesso Basin, Monticino quarry, Pellen et al., 2017). It starts with resedimented clastic evaporites (i.e. Sapigno Fm.) followed by a coarsening-and shallowingupward succession (i.e. di Tetto or San Donato Fm.) of mudstones with intercalated turbiditic sandstones (Fig. 5f) and a volcaniclastic marker bed dated initially by 40 Ar-39 Ar at ~5.5 Ma [START_REF] Odin | Integrated stratigraphy of the Maccarone section[END_REF] and later by 238 U-206 Pb at 5.5320±0.0046/0.0074 Ma [START_REF] Cosentino | Refining the Mediterranean "Messinian gap" with high-precision U-Pb zircon geochronology, central and northern Italy[END_REF]Fig. 3a). Allounit p-ev 2 (i.e. Colombacci Fm.) occurs in the deeper depocenters in 4/5 sedimentary cycles consisting of three > 5 m-thick coarse-grained bodies (conglomerates and sandstones) and two blackcolored mudstone beds alternating with fine-grained mudstones/clays with intercalated three micritic limestones (known in literature as Colombacci limestones; Figs. 3a, 5g; [START_REF] Bassetti | Paleoenvironmental significance of Messinian post-evaporitic lacustrine carbonates in the northern Apennines, Italy[END_REF]. By contrast, an incomplete Colombacci Fm. deposited in the shallower thrust-top basins (e.g. Vena del Gesso Basin and Molise sections; Pellen et al., 2017;Cosentino et al., 2018). The p-ev 2 cycles have been interpreted as reflecting the alternation of wet (mudstones and Colombacci limestones in Eastern Romagna) and dry (coarse-grained facies and Colombacci limestones in the Maccarone section) phases controlled by Milankovitch-driven climatic factors (Fig. 3a; Roveri et al., 2008a;[START_REF] Cosentino | Refining the Mediterranean "Messinian gap" with high-precision U-Pb zircon geochronology, central and northern Italy[END_REF] and, as such, they have been used for the astronomical tuning of the Colombacci Fm. to the Lago-Mare phase (Figs. 3a,4a; see subsection 2.2). By contrast, [START_REF] Clauzon | Influence of Mediterranean sea-level changes on the Dacic Basin (Eastern Paratethys) during the late Neogene: the Mediterranean Lago Mare facies deciphered[END_REF] and Popescu et al. (2007) moved the Colombacci Fm. into the Pliocene (Fig. 4b). However, this conclusion has been proven to rely on wrong stratigraphic and paleontological arguments (see Roveri et al., 2008c, Grothe et al., 2018 andsubsection 5.5). Substage 3.2 records in the Apennines do not always contain the three prominent conglomeratic facies as in Romagna, but only laminated to massive clays with sandy intercalations equivalent to the ones typifying substage 3.1 records (e.g. Maccarone section; [START_REF] Sampalmieri | Paleoenvironments of the Mediterranean Basin at the Messinian hypersaline/ hyposaline transition: evidence from natural radioactivity and microfacies of postevaporitic successions of the Adriatic sub-basin[END_REF][START_REF] Cosentino | Refining the Mediterranean "Messinian gap" with high-precision U-Pb zircon geochronology, central and northern Italy[END_REF]Fig. 5f). The absence of a lithological cyclicity that clearly mimics an orbital signal largely hampered the astronomical tuning of these claydominated sections, although an attempt has been made with the Maccarone section [START_REF] Cosentino | Refining the Mediterranean "Messinian gap" with high-precision U-Pb zircon geochronology, central and northern Italy[END_REF]. The only exception is represented by the Fonte dei Pulcini section, which has been equipped with an age framework by astronomical tuning of the magnetic susceptibility record [START_REF] Cosentino | Orbitally forced paleoenvironmental and paleoclimate changes in the late postevaporitic Messinian of the central Mediterranean Basin[END_REF]. Despite the lack of outstanding lithological changes these sections are often provided with a lithostratigraphic subdivision using the same nomenclature as in the Romagna area. When applied, the di Tetto Fm.-Colombacci Fm. boundary is placed high in the sections, i.e. few tens of meters underneath the Miocene/Pliocene boundary, resulting in a much different thickness of the formations compared to the Romagna area.

Stage 3 sediments are poorly exposed on the Tyrrhenian Sea onshore side of Italy (Fig. 2b). The best known succession crops out in the Cava Serredi quarry in the Fine Basin (Tuscany; [START_REF] Bossio | Some aspects of the upper Miocene in Tuscany[END_REF][START_REF] Bossio | Rassegna delle conoscenze sulla stratigrafia del Neoautoctono toscano[END_REF]Carnevale et al., 2006b[START_REF] Carnevale | Did the Mediterranean marine reflooding precede the Miocene/Pliocene boundary? Paleontological and geochemical evidence from upper Messinian sequences of Tuscany, Italy. Paleogeogr[END_REF]. Here the MSC has a thickness of ~150 m, of which only the uppermost ~100 m are attributed, without clear arguments, to Stage 3 by Carnevale et al. (2006b). The lowermost ~40 m of the Stage 3 unit consists of mudstone with alternating sandstone bodies which have been attributed to [START_REF] Roveri | Stratigraphy of the Messinian post-evaporitic stage in eastern-Romagna (northern Apennines, Italy)[END_REF]'s p-ev allounit, while the uppermost ~60 m form the p-ev 2 allounit and include two prominent conglomerate bodies alternating with mudstones interbedded with sandstone horizons and black, organic-rich layers (Carnevale et al., 2006b). A few and more fragmented sections are also described on the Tyrrhenian Sea side of Italy by [START_REF] Cipollari | Extension-and compression-related basins in central Italy during the Messinian Lago-Mare event[END_REF].

The Miocene/Pliocene boundary is variably expressed through the Apennine system: unconformable above the ostracod-bearing clays and highlighted by erosional surfaces draped by conglomeratic accumulations (e.g. Stingeti section in Molise; Cosentino et al., 2018), conformable above 0.5-1 m-thick black mudstones similar to how it is observed in Piedmont and of equally unknown paleoenvironmental significance (e.g. Romagna area and Maccarone section; [START_REF] Roveri | Stratigraphy of the Messinian post-evaporitic stage in eastern-Romagna (northern Apennines, Italy)[END_REF][START_REF] Gennari | The Messinian-Zanclean boundary in the Northern Apennine[END_REF] or conformable above the ostracod-rich mudstones (e.g. Maccarone and Fonte dei Pulcini sections; [START_REF] Cosentino | High-frequency cyclicity in the latest Messinian Adriatic foreland basin: insight into paleoclimate and paleoenvironments of the Mediterranean Lago-Mare episode[END_REF][START_REF] Cosentino | Orbitally forced paleoenvironmental and paleoclimate changes in the late postevaporitic Messinian of the central Mediterranean Basin[END_REF][START_REF] Cosentino | Refining the Mediterranean "Messinian gap" with high-precision U-Pb zircon geochronology, central and northern Italy[END_REF][START_REF] Sampalmieri | Paleoenvironments of the Mediterranean Basin at the Messinian hypersaline/ hyposaline transition: evidence from natural radioactivity and microfacies of postevaporitic successions of the Adriatic sub-basin[END_REF].

All p-ev 1 deposits studied are almost devoid of in-situ biota, except for fish otoliths and three fish skeletons found in the upper substage 3.1 part of Cava Serredi (Carnevale et al., 2006b). The p-ev 2 /Colombacci deposits, instead, host typical Paratethyan assemblages of brackishwater mollusks, ostracods, dinocysts and fish [START_REF] Bassetti | Ostracod faunas and brackish-water environments of the late Messinian Sapigno section (northern Apennines, Italy)[END_REF][START_REF] Bertini | The Northern Apennines palynological record as a contribute for the reconstruction of the Messinian paleoenvironments[END_REF]Popescu et al., 2007;[START_REF] Grossi | Late Messinian Lago-Mare ostracods and paleoenvironments of the central and eastern Mediterranean Basin[END_REF][START_REF] Iaccarino | The Trave section (Monte dei Corvi, Ancona, central Italy): an integrated paleontological study of the Messinian deposits[END_REF][START_REF] Cosentino | Orbitally forced paleoenvironmental and paleoclimate changes in the late postevaporitic Messinian of the central Mediterranean Basin[END_REF]Cosentino et al., , 2018;;[START_REF] Schwarzhans | Late Miocene-Early Pliocene evolution of Mediterranean gobies and their environmental and biogeographic significance[END_REF]. A diverse array of marine fossils (benthic and planktic foraminifera, calcareous nannofossils, dinocysts and fish otoliths and skeletons) has also been reported from the horizons containing these Paratethyan taxa [START_REF] Bertini | The Northern Apennines palynological record as a contribute for the reconstruction of the Messinian paleoenvironments[END_REF]Carnevale et al., 2006a;Popescu et al., 2007;Pellen et al., 2017). While the autochthony of ostracods, when considered, is unquestioned, the allochthonous vs autochthonous character of the other mentioned fossils is disputed and still unclear (see Chapter 5).

Sicily

The MSC record is widely exposed on Sicily, mainly in the Caltanissetta Basin and in scattered locations on the Hyblean Plateau (i.e. Ragusa-Siracusa area) and the Messina area (Fig. 2b; [START_REF] Butler | Tectonics and sequence stratigraphy in Messinian Basins, Sicily: constraints on the initiation and termination of the Mediterranean salinity crisis[END_REF][START_REF] Manzi | A new facies model for the Upper Gypsum of Sicily (Italy): chronological and paleoenvironmental constraits for the Messinian salinity crisis in the Mediterranean[END_REF][START_REF] Sciuto | New reports of Paratethyan ostracods affinity from the Mediterranean Basin (Sicily, Italy)[END_REF]. Like the Northern Apennines, it shows a complex distribution and variable stratigraphy that mirrors the structuring of Sicily into basins with different characters, geometries and depocenters which subsided at different times and rates [START_REF] Butler | Tectonics and sequence stratigraphy in Messinian Basins, Sicily: constraints on the initiation and termination of the Mediterranean salinity crisis[END_REF][START_REF] Catalano | Sicily's fold-thrust belt and slab roll-back: the SI. RI. PRO. seismic crustal transect[END_REF]. This structural setting permitted the simultaneous deposition of shallow and intermediate-water sediments (Roveri et al., 2008b). Mostly found in the Caltanissetta Basin, these intermediate-water successions have for decades been considered the onshore counterpart of the offshore evaporitic trilogy seen in seismic data from the Western Mediterranean Basin [START_REF] Decima | Late Miocene evaporites of the Central Sicilian Basin[END_REF]. More recently, Raad et al. (2021) attempted a similar onshore-offshore correlation but with the intermediate Central Mallorca Depression. The currently endorsed stratigraphic model (Fig. 4a), refined over the years by [START_REF] Decima | Ostracodi del Gen. Cyprideis JONES del Neogene e del Quaternario italiani[END_REF]Wezel (1971, 1973), [START_REF] Decima | The origin of "evaporative" limestones: an example from the Messinian of Sicily (Italy)[END_REF] Roveri et al. (2008b) and [START_REF] Manzi | A new facies model for the Upper Gypsum of Sicily (Italy): chronological and paleoenvironmental constraits for the Messinian salinity crisis in the Mediterranean[END_REF], envisages two 'evaporitic cycles'. The 'First cycle', overlying both alluvial and deep-water sediments (Tripoli Fm., Licata Fm. and Terravecchia Fm.;see Maniscalco et al., 2019 and references therein), comprises the disputed Calcare di Base [START_REF] Manzi | The Messinian CdB (Sicily, Italy) revisited[END_REF], 2016bvs Caruso et al., 2015), PLG or Gessi di Cattolica Fm. [START_REF] Decima | Late Miocene evaporites of the Central Sicilian Basin[END_REF]Lugli et al., 2010) and the Halite Unit [START_REF] Lugli | Giant polygons in the Realmonte mine (Agrigento, Sicily): evidence for the desiccation of a Messinian halite basin[END_REF]. The 'Second cycle' comprises the Upper Gypsum (UG) or Gessi di Pasquasia Fm., which is only present in depocenters of the Caltanissetta Basin (see [START_REF] Manzi | A new facies model for the Upper Gypsum of Sicily (Italy): chronological and paleoenvironmental constraits for the Messinian salinity crisis in the Mediterranean[END_REF] for a detailed overview), sporadically overlain by the siliciclastic Arenazzolo Fm. [START_REF] Decima | Late Miocene evaporites of the Central Sicilian Basin[END_REF][START_REF] Cita | Sedimentation in the latest Messinian at Capo Rossello (Sicily)[END_REF]. The whole succession is sealed by the Pliocene marine Trubi Fm. (Fig. 4a). The two evaporite cycles are separated by an erosional surface (MES) associated with an angular discordance broadly linked to the main Mediterranean drawdown event (e.g. [START_REF] Butler | Tectonics and sequence stratigraphy in Messinian Basins, Sicily: constraints on the initiation and termination of the Mediterranean salinity crisis[END_REF]Roveri et al., 2008b). [START_REF] Andreetto | Alternate interpretation of the Messinian salinity crisis, controversy resolved?[END_REF], however, placed the MES at the Arenazzolo Fm.-Trubi Fm. transition, implying that the entire evaporitic deposition in the Caltanissetta Basin pre-dated the offshore one, but they do not provide evidence of erosion at that level. In more recent publications from the same research group, the MES is shifted towards the base of the Arenazzolo Fm. (e.g. [START_REF] Bache | A twostep process for the reflooding of the Mediterranean after the Messinian Salinity Crisis[END_REF], again without evidence of major erosion, and different ages are assigned (see Fig. 4b and Grothe et al., 2018 for details).

APPENDIX D

The Upper Gypsum successions are commonly incomplete in many of the Caltanissetta Basin sections (Pasquasia-Capodarso, Casteltermini, Alimena, Nicosia, Siculiana-Marina; [START_REF] Decima | Comments on late Messinian microfaunas in several sections from Sicily[END_REF]Rouchy and Caruso, 2006;[START_REF] Manzi | A new facies model for the Upper Gypsum of Sicily (Italy): chronological and paleoenvironmental constraits for the Messinian salinity crisis in the Mediterranean[END_REF]Fig. 5i). In the most complete section, Eraclea Minoa (Fig. 3a), the Upper Gypsum Unit consists of 6 (Van der Laan et al., 2006) to 7 [START_REF] Manzi | A new facies model for the Upper Gypsum of Sicily (Italy): chronological and paleoenvironmental constraits for the Messinian salinity crisis in the Mediterranean[END_REF] primary gypsum beds with a repetitive internal organization of facies (see [START_REF] Schreiber | Field trip to Eraclea Minoa: Upper Messinian. "Neogene Mediterranean Paleoceanography[END_REF][START_REF] Manzi | A new facies model for the Upper Gypsum of Sicily (Italy): chronological and paleoenvironmental constraits for the Messinian salinity crisis in the Mediterranean[END_REF] for facies description) interbedded with marls and lenticular terrigenous sandstone bodies, gypsarenites and gypsrudites (Fig. 5h). Two of the terrigenous sandstone bodies are highlighted by [START_REF] Manzi | A new facies model for the Upper Gypsum of Sicily (Italy): chronological and paleoenvironmental constraits for the Messinian salinity crisis in the Mediterranean[END_REF] in the thick (~60 m), Cyprideis agrigentina-rich (Grossi et al., 2015), marly interval dividing gypsum VI and VII for its alleged astronomical significance (Fig. 3a; see subsection 2.2). A mixed (physically reworked) marine (foraminifera and dinocysts) and (in-situ) brackish biota (ostracods and dinocysts) of Paratethyan origin characterizes the marly interbeds from at least gypsum III upwards (following the investigations carried on the Eraclea Minoa section; [START_REF] Bonaduce | Paleoecological interpretation of the latest Messinian sediments from southern Sicily (Italy)[END_REF]Rouchy and Caruso, 2006;[START_REF] Londeix | Messinian paleoenvironments and hydrology in Sicily (Italy): the dinoflagellate cyst record[END_REF]Grossi et al., 2015;Fig. 3a). Calcareous nannofossils have been found along with the above organisms in a more northerly location by Maniscalco et al. (2019) and considered reworked. Above the last gypsum, the ~6-7 m-thick Arenazzolo Fm. is found, represented by reddish arkosic cross-laminated and poorly consolidated sand [START_REF] Bonaduce | Paleoecological interpretation of the latest Messinian sediments from southern Sicily (Italy)[END_REF]Roveri et al., 2008b) and interpreted as the expression of a shallow-water delta, albeit without a sedimentological investigation (e.g. [START_REF] Decima | Late Miocene evaporites of the Central Sicilian Basin[END_REF][START_REF] Cita | Sedimentation in the latest Messinian at Capo Rossello (Sicily)[END_REF]. The whole Stage 3 sequence is conformably overlain by the Zanclean marine Trubi Fm. in the basin center (e.g. at Eraclea Minoa and Capo Rossello;Fig. 6d;[START_REF] Brolsma | Lithostratigraphy and Foraminiferal assemblages of the Miocene/ Pliocene transitional strata at Capo Rossello and Eraclea Minoa (Sicily, Italy): Kon[END_REF][START_REF] Cita | Sedimentation in the latest Messinian at Capo Rossello (Sicily)[END_REF][START_REF] Van Couvering | The base of the Zanclean Stage and of the Pliocene Series[END_REF]Rouchy and Caruso, 2006;[START_REF] Manzi | A new facies model for the Upper Gypsum of Sicily (Italy): chronological and paleoenvironmental constraits for the Messinian salinity crisis in the Mediterranean[END_REF]Fig. 6d) and unconformably in the shallower, marginal areas [START_REF] Manzi | A new facies model for the Upper Gypsum of Sicily (Italy): chronological and paleoenvironmental constraits for the Messinian salinity crisis in the Mediterranean[END_REF]Roveri et al., 2019b). Only [START_REF] Decima | Late Miocene evaporites of the Central Sicilian Basin[END_REF] and Raad et al. (2021) report the Miocene/Pliocene transition in the key, intermediate water-representative section of Eraclea Minoa as erosive. However, they do not provide evidence (e.g. photographic documentation) for the presence of an erosional unconformity and, moreover, Raad et al. (2021) erroneously refer to [START_REF] Cita | Sedimentation in the latest Messinian at Capo Rossello (Sicily)[END_REF], where no erosion is mentioned at the M/P boundary in Eraclea Minoa.

The bathymetric jump between the <100 m of water depth during the late Messinian and the >200 m at the base of the Trubi Fm. is often regarded as a key onshore evidence of the sudden and catastrophic Mediterranean-Atlantic re-connection at the Miocene/Pliocene boundary (e.g. [START_REF] Caruso | The late Messinian "Lago-Mare" event and the Zanclean Reflooding in the Mediterranean Sea: new insights from the Cuevas del Almanzora section (Vera Basin, South-Eastern Spain)[END_REF]. However, the real depth of the base of the Trubi is all but obvious. In fact, variable estimates have been proposed based on the observed benthic foraminifera and/or psychrospheric ostracods at Capo Rossello and Eraclea Minoa: 200-500 m [START_REF] Decima | Late Miocene evaporites of the Central Sicilian Basin[END_REF], 600-800 m [START_REF] Sgarrella | Paleoceanographic conditions at the base of the Pliocene in the Southern Mediterranean Basin[END_REF][START_REF] Sgarrella | The Capo Rossello Bore-Hole (Agrigento, Sicily): cyclostratigraphic and F. Andreetto et al. paleoceanographic reconstructions from quantitative analyses of the Zanclean foraminiferal assemblages[END_REF][START_REF] Barra | Paleoenvironmental bottom water conditions in the early Zanclean of the Capo Rossello area (Agrigento, Sicily)[END_REF], 1400-2400 m [START_REF] Cita | Sedimentation in the latest Messinian at Capo Rossello (Sicily)[END_REF].

Greece

Several MSC localities are reported from the Greek Ionian Islands (Corfu, Cephalonia and Zakynthos) and from Crete (Fig. 2b).

On the NW coast of Corfu (Aghios Stefanos section), the PLG unit is missing and only a 32 m-thick cyclically-arranged terrigenous succession is present comprising three m-thick conglomerate beds alternating with fine-grained deposits rich in unspecified species of brackish water ostracods [START_REF] Pierre | Reconstruction of the paleoenvironmental changes around the Miocene-Pliocene boundary along a West-East transect across the Mediterranean[END_REF].

In the southern part of Zakynthos, an evaporitic succession composed of eight gypsum cycles (Kalamaki section) occurs above marine marly deposits (Karakitsios et al., 2017b). These gypsum beds were initially ascribed to the UG unit [START_REF] Pierre | Reconstruction of the paleoenvironmental changes around the Miocene-Pliocene boundary along a West-East transect across the Mediterranean[END_REF] and later to the PLG (Karakitsios et al., 2017b). The gypsum unit is overlain by approximately ~13 m of siltstones and marls with scattered, cm-thick beds of sandstones, conglomerates and carbonates with nodular texture [START_REF] Pierre | Reconstruction of the paleoenvironmental changes around the Miocene-Pliocene boundary along a West-East transect across the Mediterranean[END_REF]Karakitsios et al., 2017b). Although no ostracods are reported from this interval, due to its stratigraphic position the postevaporitic unit is correlated to the Lago-Mare phase (Karakitsios et al., 2017b). Except for the rare presence of marine nannofossils (Ceratolithus acutus together with Reticulofenestra zancleana) just below the Miocene/ Pliocene boundary, only reworked marine fauna has been reported from the post-evaporitic package (Karakitsios et al., 2017b). This dominantly terrigenous succession is unconformably overlain by the Zanclean Trubi Formation (Fig. 6e; Karakitsios et al., 2017b).

MSC deposits on Crete (e.g. [START_REF] Meulenkamp | Field Guide to the Neogene of Crete[END_REF][START_REF] Delrieu | La surface d'erosion finimessinienne en Crète centrale (Grèce) et sur le pourtour méditerranéen: Rapports avec la crise de salinité méditerranéenne[END_REF][START_REF] Cosentino | The late Messinian Lago-Mare episode in the Mediterranean Basin: preliminary report on the occurrence of Paratethyan ostracod fauna from central Crete (Greece)[END_REF]Roveri et al., 2008a;[START_REF] Zachariasse | Mass wasting and uplift on Crete and Karpathos (Greece) during the Early Pliocene related to beginning of South Aegean left-lateral, strike slip tectonics[END_REF][START_REF] Zachariasse | Formation and fragmentation of a late Miocene supradetachment basin in central Crete: implications for exhumation mechanisms of high-pressure rocks in the Aegean forearc[END_REF] were deposited in Miocene extensional, fault-bound basins driven by tectonic subsidence that ceased in the late Pliocene and Pleistocene [START_REF] Van Hinsbergen | Neogene supradetachment basin development on Crete (Greece) during exhumation of the South Aegean core complex[END_REF]. Because of the strong tectonic and eustatic sea-level-related fragmentation of the stratigraphic record, reconstructing the late Miocene stratigraphy of Crete has not been straightforward [START_REF] Zachariasse | Mass wasting and uplift on Crete and Karpathos (Greece) during the Early Pliocene related to beginning of South Aegean left-lateral, strike slip tectonics[END_REF][START_REF] Zachariasse | Formation and fragmentation of a late Miocene supradetachment basin in central Crete: implications for exhumation mechanisms of high-pressure rocks in the Aegean forearc[END_REF]. Several primary and clastic gypsum facies are recognized, but their correlation with the MSC stratigraphy is disputed (see [START_REF] Cosentino | The late Messinian Lago-Mare episode in the Mediterranean Basin: preliminary report on the occurrence of Paratethyan ostracod fauna from central Crete (Greece)[END_REF]Roveri et al., 2008aRoveri et al., , 2014a;;[START_REF] Zachariasse | Mass wasting and uplift on Crete and Karpathos (Greece) during the Early Pliocene related to beginning of South Aegean left-lateral, strike slip tectonics[END_REF]. Coarsegrained, mammal-bearing terrigenous facies irregularly alternating with marls are in places found unconformably overlying the gypsum and separated from the Pliocene facies by an erosion surface (see [START_REF] Meulenkamp | Field Guide to the Neogene of Crete[END_REF][START_REF] Delrieu | La surface d'erosion finimessinienne en Crète centrale (Grèce) et sur le pourtour méditerranéen: Rapports avec la crise de salinité méditerranéenne[END_REF][START_REF] Cosentino | The late Messinian Lago-Mare episode in the Mediterranean Basin: preliminary report on the occurrence of Paratethyan ostracod fauna from central Crete (Greece)[END_REF]. In two localities on the Messarà Plain, [START_REF] Cosentino | The late Messinian Lago-Mare episode in the Mediterranean Basin: preliminary report on the occurrence of Paratethyan ostracod fauna from central Crete (Greece)[END_REF] described a highly diversified ostracod fauna with Paratethyan affinity in some marly intervals.

Messinian evaporites and/or Lago-Mare deposits are also reported from the North Aegean region onshore in the Strymon Basin [START_REF] Snel | Calcareous nannofossil biostratigraphy and magnetostratigraphy of the upper Miocene and lower Pliocene of the Northern Aegean (Orphanic Gulf-Strimon Basin areas), Greece. Paleogeogr[END_REF][START_REF] Suc | Marine gateway vs. fluvial stream within the Balkans from 6 to 5 Ma[END_REF]Karakitsios et al., 2017a) and Dardanelles region [START_REF] Melinte-Dobrinescu | The Messinian salinity crisis in the Dardanelles region: Chronostratigraphic constraints[END_REF] and offshore (Prinos-Nestos Basin; Karakitsios et al., 2017a), but recent integrated studies suggested that the sections studied by the above listed authors are older than the MSC (see Krijgsman et al., 2020aKrijgsman et al., , 2020b)). In particular, Krijgsman et al. (2020a) proposed that for most, if not all, of the MSC the North Aegean was a brackish water, mostly Paratethyan-fed basin restricted by the Cyclades sill to the south (Fig. 2a) and forming a passageway for Paratethyan overspill waters towards the Mediterranean.

Cyprus

MSC deposits on Cyprus outcrop in the Pissouri, Psematismenos, Mesaoria and Polemi basins on the southerly fringe of the Troodos massif (Fig. 2b; [START_REF] Rouchy | Paleoenvironmental changes at the Messinian-Pliocene boundary in the eastern Mediterranean (southern Cyprus basins): significance of the Messinian Lago-Mare[END_REF]Manzi et al., 2016a). According to F. Andreetto et al. [START_REF] Rouchy | Paleoenvironmental changes at the Messinian-Pliocene boundary in the eastern Mediterranean (southern Cyprus basins): significance of the Messinian Lago-Mare[END_REF] and [START_REF] Orszag-Sperber | The onset of the Messinian salinity crisis: insights from Cyprus sections[END_REF], sediments belonging to all MSC stages of Roveri et al. (2014a) are preserved in the Cypriot basins. By contrast, [START_REF] Robertson | Depositional processes and basin analysis of Messinian evaporites in Cyprus[END_REF] and Manzi et al. (2016a) considered that PLG evaporites on Cyprus are only present as fragments reworked within a chaotic unit (the Lower Gypsum and Intermediate breccia units of [START_REF] Orszag-Sperber | The onset of the Messinian salinity crisis: insights from Cyprus sections[END_REF] and that the only in situ evaporites belong to the overlying Upper Gypsum Unit, which encompasses the whole of Stage 3 (Figs. 3a,4a). A continuous, Eraclea Minoa-like section is not known in Cyprus (Manzi et al., 2016a). The best exposure of the lower 60 m of this unit is found in the Polemi Basin (Manzi et al., 2016a). It comprises up to six gypsum beds (the lower three of which are mainly selenitic, while the upper three are predominantly laminated; Fig. 3a). Gypsum beds range in thickness from 1 to 6 m and are separated by laminated marls (Fig. 5j) occasionally interbedded with conglomerates and sandstones (e.g. between the 5 th and 6 th gypsum layers; [START_REF] Rouchy | La crise evaporitique messinienne de Mediterranee: nouvelles propositions pour une interpretation genetique[END_REF][START_REF] Rouchy | Paleoenvironmental changes at the Messinian-Pliocene boundary in the eastern Mediterranean (southern Cyprus basins): significance of the Messinian Lago-Mare[END_REF]Manzi et al., 2016a). The sixth gypsum bed is reported by [START_REF] Rouchy | Paleoenvironmental changes at the Messinian-Pliocene boundary in the eastern Mediterranean (southern Cyprus basins): significance of the Messinian Lago-Mare[END_REF] to be hollowed in the upper part with the cavities filled with overlying sediments. The similarity of the cyclicity and facies association of this Cyprus succession with the substage 3.1 interval of the Sicilian UG led Manzi et al. (2016a) to propose a bed-to-bed correlation and to recognize the substage 3.1/ 3.2 boundary at the top of the last gypsum bed (Fig. 3a). According to [START_REF] Orszag-Sperber | La transition Messinien-Pliocène en Méditerranée orientale (Chypre): la période du Lago-Mare et sa signification[END_REF] and [START_REF] Rouchy | Paleoenvironmental changes at the Messinian-Pliocene boundary in the eastern Mediterranean (southern Cyprus basins): significance of the Messinian Lago-Mare[END_REF], this chronostratigraphic boundary coincides with a Mediterranean-scale sealevel drop, a conclusion that arises from the interpretation of the cavities in the uppermost gypsum as the product of karstic dissolution following a prolonged period of subaerial exposure.

The sedimentary sequence overlying the last gypsum bed and assigned by Manzi et al. (2016a) to the Lago-Mare phase lacks a clear and rhythmic sedimentary cyclicity. In the Pissouri Basin this interval (up to 25-30 m-thick) mostly consists of conglomerates, sandstones, limestones, paleosols (which appear as dm to m-thick dark marly horizons, in one case with pulmonated gastropods) and subordinated claymarly horizons [START_REF] Rouchy | Paleoenvironmental changes at the Messinian-Pliocene boundary in the eastern Mediterranean (southern Cyprus basins): significance of the Messinian Lago-Mare[END_REF]. By contrast, in the Polemi sections the clay-marly facies dominates this interval [START_REF] Rouchy | Paleoenvironmental changes at the Messinian-Pliocene boundary in the eastern Mediterranean (southern Cyprus basins): significance of the Messinian Lago-Mare[END_REF]. In situ fresh-brackish water species of articulated mollusks (Limnocardiidae, Melanopsis), Paratethyan (Loxocorniculina djafarovi, Euxynocythere praebaquana) and Mediterranean (Cyprideis agrigentina) ostracods and foraminifera (Ammonia beccarii), Characeae, abundant fragments of the marine euryhaline fish Clupeidae and a fish skeleton of the euryhaline Aphanius crassicaudus are described from some of the substage 3.1 and 3.2 fine-grained facies and within the terrigenous laminae of some balatino gypsum [START_REF] Orszag-Sperber | La transition Messinien-Pliocène en Méditerranée orientale (Chypre): la période du Lago-Mare et sa signification[END_REF][START_REF] Rouchy | Paleoenvironmental changes at the Messinian-Pliocene boundary in the eastern Mediterranean (southern Cyprus basins): significance of the Messinian Lago-Mare[END_REF][START_REF] Orszag-Sperber | Changing perspectives in the concept of "Lago-Mare" in Mediterranean Late Miocene evolution[END_REF]Manzi et al., 2016a). The upward change in diversity of the ostracod fauna seen elsewhere (e.g. Malaga, Nijar, Vera, Apennines and Eraclea Minoa) is not reported in Cyprus but this may be because no detailed study of ostracod assemblages in Stage 3 sediments has been published. The Miocene/Pliocene boundary, near Polemi village is described by Manzi et al. (2016a) as a sharp contact above a dark, organic-rich layer (Fig. 6f). It appears to be similar to the boundary reported from Piedmont (Fig. 6c; [START_REF] Trenkwalder | The Miocene/Pliocene boundary in the Early Pliocene micropaleontological record: new data from the Tertiary Piedmont Basin (Moncucco quarry, Torino Hill, northwestern Italy)[END_REF][START_REF] Pierre | The northermost record of the Messinian salinity crisis (Piedmont Basin, NW Italy)[END_REF] and Northern Apennines [START_REF] Gennari | The Messinian-Zanclean boundary in the Northern Apennine[END_REF][START_REF] Grossi | Late Messinian Lago-Mare ostracods and paleoenvironments of the central and eastern Mediterranean Basin[END_REF], if not for the presence, in Cyprus, of (possibly) insitu Cyprideis agrigentina (Manzi et al., 2016a). A layer with the same field appearance, thickness (~ 1 m) and stratigraphic position is reported in Pissouri by [START_REF] Rouchy | Paleoenvironmental changes at the Messinian-Pliocene boundary in the eastern Mediterranean (southern Cyprus basins): significance of the Messinian Lago-Mare[END_REF], which they interpreted as a paleosol based on mottling, oxidized roots, carbonate concretions and plant fragments.

Southern Turkey

The tectonically active, during the Miocene, thrust-top basin of Adana in southern Turkey [START_REF] Radeff | Sedimentary evidence for late Messinian uplift of the SE margin of the Central Anatolian Plateau: Adana Basin, southern Turkey[END_REF] retains the most complete and better exposed easternmost successions of the MSC (Fig. 2b), whose deposits were attributed to the Handere Fm. [START_REF] Cosentino | The Messinian salinity crisis in the marginal basins of the peri-Mediterranean orogenic systems: examples from the central Apennines (Italy) and the Adana Basin (Turkey)[END_REF][START_REF] Radeff | Stratigraphic architecture of the upper Messinian deposits of the Adana Basin (southern Turkey): implications for the Messinian salinity crisis and the Taurus petroleum system[END_REF].

MSC Stage 3 finds expression in a >1 km thick continental unit unconformable, through an erosional surface, above the pre-evaporitic, Stage 1 anhydrite-shale alternations [START_REF] Radeff | Stratigraphic architecture of the upper Messinian deposits of the Adana Basin (southern Turkey): implications for the Messinian salinity crisis and the Taurus petroleum system[END_REF] and resedimented gypsum-bearing Stage 2 deposits [START_REF] Cosentino | The Messinian salinity crisis in the marginal basins of the peri-Mediterranean orogenic systems: examples from the central Apennines (Italy) and the Adana Basin (Turkey)[END_REF][START_REF] Cipollari | Easternmost Mediterranean evidence of the Zanclean flooding event and subsequent surface uplift: Adana Basin, southern Turkey[END_REF]. This unit mainly consists of fluvial coarse-and fine-grained deposits representing channel fill and overbank deposits. Sporadically, some fine-grained intercalations are found containing a mixed brackish (ostracod) and marine (foraminifera and calcareous nannofossils) fauna. The ostracod fauna has affinity with the Paratethyan fauna but, unlike to many other Mediterranean onshore localities, is poorly diversified, with monospecific assemblages of Cyprideis agrigentina (Avadan section and T-191 borehole; [START_REF] Cipollari | Easternmost Mediterranean evidence of the Zanclean flooding event and subsequent surface uplift: Adana Basin, southern Turkey[END_REF] or with Cyprideis agrigentina accompanied by rare to abundant specimens of Loxoconcha muelleri, Euxinocyhere (Maeotocythere) praebaquana, and Loxoconcha sp. (Adana section; [START_REF] Faranda | Messinian paleoenvironmental changes in the easternmost Mediterranean Basin: Adana Basin, southern Turkey[END_REF]. Ostracods are often associated with Ammonia beccarii and rare Elphidium and Cribroelphidium, which are the only foraminifera considered as autochthonous. Conversely, the entire nannoflora is interpreted as physically reworked [START_REF] Cipollari | Easternmost Mediterranean evidence of the Zanclean flooding event and subsequent surface uplift: Adana Basin, southern Turkey[END_REF][START_REF] Faranda | Messinian paleoenvironmental changes in the easternmost Mediterranean Basin: Adana Basin, southern Turkey[END_REF].

The Handere Fm. is followed by early Zanclean marine sediments (Avadan Fm.) deposited, according to the paleoecology of the benthic foraminifera species recognized, at bathymetries ranging from 200 to 500 m [START_REF] Cipollari | Easternmost Mediterranean evidence of the Zanclean flooding event and subsequent surface uplift: Adana Basin, southern Turkey[END_REF]. The lithological nature of the Miocene/ Pliocene boundary in the Adana Basin is not clear, but it occurs either above the continental or subaqueous, ostracod-bearing facies.

A similar stratigraphic sequence is present in the subsurface. Here, however, chaotic gypsum-bearing deposits are not found and two halite bodies ~20 and ~170 m-thick are present, separated and followed by fluvial gravels, sands and silts [START_REF] Cipollari | Easternmost Mediterranean evidence of the Zanclean flooding event and subsequent surface uplift: Adana Basin, southern Turkey[END_REF].

Summary of the onshore Stage 3 record

Most of the onshore Stage 3 records formed in shallow marginal Mediterranean basins, which underwent substantial uplift from the Messinian till nowadays and are assumed to have had their depocenter at ~200 to 50 m below the Atlantic level during the late Messinian (Roveri et al., 2014a(Roveri et al., , 2019a;;[START_REF] Radeff | Stratigraphic architecture of the upper Messinian deposits of the Adana Basin (southern Turkey): implications for the Messinian salinity crisis and the Taurus petroleum system[END_REF][START_REF] Radeff | Sedimentary evidence for late Messinian uplift of the SE margin of the Central Anatolian Plateau: Adana Basin, southern Turkey[END_REF]. The Caltanissetta Basin (Sicily), some basins along the Apennines and (possibly) Cyprus represent, instead, possible onshore representative of intermediate basins. The nature and duration of these records is quite variable, and there are only six sections that may record an entire Stage sequence (i.e. Sorbas, Nijar, Northern Apennines, Eraclea Minoa and Cyprus; Fig. 3a). Reasons for the fragmentary nature of the Stage sedimentary record include different durations of subaerial exposure following the Stage 2 drawdown, local tectonics and associated syndepositional erosion and deposition. One of the consequences of this is that any sedimentary cyclicity that resulted from orbital fluctuations is typically either less well developed or poorly preserved, making the chronology of Stage 3 rather uncertain or controversial in places.

Despite this variability, several fairly consistent characteristics are widely expressed. These are: sediments of Lago-Mare successions. Only in the intermediate Caltanissetta Basin (Sicily) do these diversified Paratethyan forms (only ostracods) occur earlier, in the sediments from substage 3.1. Where these have been studied in detail, these assemblages typically show an increase in diversity with time (possibly every wet phase of the precession cycles). Some of these sediments also contain marine fossils and there is controversy over whether these are in situ and contemporaneous or reworked. 4) In outcrop, the Miocene/Pliocene boundary has four main sedimentary expressions: erosive and followed by a conglomeratic lag (e. g. Malaga, Vera, Mallorca; Fig. 6a); conformable above continental facies (e.g. Nijar Basin; Fig. 6b); conformable above ostracod-rich mudstones (e.g. Eraclea Minoa; Fig. 6d); sharp contact above a black layer of still largely unknown paleoenvironmental significance (Piedmont, Apennines and Cyprus; Fig. 6c,f).

For a better understanding of how Stage 3 developed across the Mediterranean these marginal records now need to be considered alongside the evidence from intermediate to deep offshore settings.

We note that alternative chronostratigraphic frameworks have been proposed for several onshore (Malaga, Sorbas, Mallorca, Apennines, Sicily) and offshore (Sites 134B, 976B, 978A) locations (see Fig. 4b for references), but we have omitted them as they are shown to rely on incorrect (bio)stratigraphic arguments (see Roveri et al., 2008c, Grothe et al., 2018 andsubsection 5.5).

Offshore domain

The offshore Mediterranean is a complex array of variable-depth and morphologically complex subbasins framed by morphological highs or sills. Traditionally it is divided into two main domains (Fig. 2a), the Western and Eastern Mediterranean, with the intervening divide (or Sicily sill) situated in the Sicily channel at present with a maximum depth of 316 m. The Alborán Basin, the depressions on the Balearic Promontory, the Corsica, Valencia, Algero-Balearic, Liguro (or Sardo)-Provençal and Tyrrhenian basins belong to the "Western Mediterranean" (Fig. 2a). The Adriatic foredeep, the Ionian, Sirte, Aegean and Levant basins belong in the "Eastern Mediterranean" (Fig. 2a). Smaller-sized depressions, again surrounded by sills of variable depth, are identified and labelled within each of these subbasins.

Although the exact topography and hypsometry of the Messinian Mediterranean is difficult to reconstruct, this present-day geography is generally assumed to have been in place, with five main differences: (1) the Alborán Basin was split into a Western (WAB) and Eastern Alborán (EAB) by a volcanic arc (Booth-Rea et al., 2018); (2) the Tyrrhenian Basin was only partly opened [START_REF] Lymer | The Western Tyrrhenian Sea revisited: New evidence for a rifted basin during the Messinian Salinity Crisis[END_REF]; (3) the precise depth and width of the ancient Sicily Sill are difficult to estimate, but may have been much deeper than today (~300 m; Meijer and Krijgsman, 2005Jolivet et al., 2006). Paleodepth estimations for the Messinian configuration range from 380 m [START_REF] Just | Erosion of continental margins in the Western Mediterranean due to sea-level stagnancy during the Messinian Salinity Crisis[END_REF] to 430 m [START_REF] Garcia-Castellanos | Catastrophic flood of the Mediterranean after the Messinian salinity crisis[END_REF]; (4) one or two sills were present at the southern termination of the Adriatic foredeep (Pellen et al., 2017;[START_REF] Amadori | Restored topography of the Po Plain-Northern Adriatic region during the Messinian base level drop-Implications for the physiography and compartmentalization of the paleo-Mediterranean basin[END_REF]Manzi et al., 2020); [START_REF] Cita | Inventory of biostratigraphical findings and problems[END_REF] the North Aegean was partially isolated from the Mediterranean by the Cyclades Sill and with high Paratethys affinity (Krijgsman et al., 2020a). Following the schematic classification of the Messinian sub-basins by Roveri et al. (2014a), all these subbasins are regarded as either intermediate (i.e. relatively deep-water, 200-1000 m) or deep (water depth > 1000 m).

Compared with the onshore domain, the offshore basins hold a far greater percentage of the total volume of MSC sediments (Ryan, 1973;Haq et al., 2020). The architecture, geometry and main lithologies of the MSC and younger deposits are well known thanks to the high density of seismic data and the fact that evaporites (halite particularly) are easily identified on seismic profiles due to their unusual seismic properties, especially compared to those of terrigenous and carbonate sediments (see Lofi et al., 2011aLofi et al., , 2011b;;Lofi, 2018;Haq et al., 2020). However, the detailed lithological, sedimentological, paleontological and geochemical nature and their chronostratigraphy are still poorly constrained offshore because these cannot be univocally defined on the basis of seismic data alone (Roveri et al., 2019b) and direct information about these deep MSC successions is limited to a small number of cores ( 16) drilled during the DSDP (Ryan et al., 1973;[START_REF] Hsü | Initial reports of the deep sea drilling project: DSDP volume XLII Part[END_REF] and ODP [START_REF] Kastens | Proc. ODP, Init. Repts[END_REF][START_REF] Comas | Proceedings of the Ocean Drilling Program, Initial Reports[END_REF][START_REF] Emeis | Proc. ODP, Initial Reports[END_REF] drilling campaigns that penetrated exclusively the uppermost tens of meters of the deep MSC suite in very scattered localities (Fig. 2b). Only recently, access to industrial boreholes crossing the base of the halite in the deep Levant Basin has been granted, providing a rare glimpse of the deep MSC deposits in the easternmost part of the Mediterranean [START_REF] Gvirtzman | Intra-Messinian truncation surface in the Levant Basin explained by subaqueous dissolution[END_REF][START_REF] Manzi | The onset of the Messinian salinity crisis in the deep Eastern Mediterranean basin[END_REF][START_REF] Meilijson | Deep-basin evidence resolves a 50-year-old debate and demonstrates synchronous onset of Messinian evaporite deposition in a non-desiccated Mediterranean[END_REF]Meilijson et al., , 2019)).

The MSC is commonly described as tripartite ('Messinian trilogy' after [START_REF] Montadert | Messinian event: seismic evidence[END_REF] in the Western Mediterranean (Lower-Mobile-Upper units: LU-MU-UU, respectively). However, in the Ionian Basin is described as bipartite (MU-UU) by [START_REF] Camerlenghi | Seismic markers of the Messinian salinity crisis in the deep Ionian Basin[END_REF] while according to Lofi et al. (2011a), [START_REF] Gvirtzman | Intense salt deformation in the Levant Basin in the middle of the Messinian salinity crisis[END_REF][START_REF] Gvirtzman | Intra-Messinian truncation surface in the Levant Basin explained by subaqueous dissolution[END_REF], Lofi (2018) and [START_REF] Camerlenghi | Seismic markers of the Messinian salinity crisis in the deep Ionian Basin[END_REF], the Levant Basin consists of the MU and the UU is only present locally and possibly represented by evaporite-free terrigenous accumulations [START_REF] Kartveit | Evidence of sea level drawdown at the end of the Messinian salinity crisis and seismic investigation of the Nahr Menashe unit in the northern Levant Basin, offshore Lebanon[END_REF][START_REF] Madof | Discovery of vast fluvial deposits provides evidence for drawdown during the late Miocene Messinian salinity crisis[END_REF]. The lack of many age constraints within the offshore MSC successions hampers unambiguous correlation with onshore sequences (Fig. 1a; Roveri et al., 2014a). Nevertheless, different authors have proposed onshore-offshore correlation of specific events (e.g. onset, Ochoa et al., 2015; and termination of the MSC, [START_REF] Biscaye | Age and nature of the Pan-Mediterranean subbottom reflector M. The Mediterranean Sea[END_REF], Iaccarino et al., 1999) and stratigraphic schemes [START_REF] Decima | Osservazioni sulle evaporiti Messiniane della Sicilia centro-meridionale[END_REF]Raad et al., 2021) based on and biostratigraphic evidence [START_REF] Cosentino | Environments and tectonic instability in central Italy (Garigliano Basin) during the late Messinian Lago-Mare episode: new data from the onshore Mondragone 1 well[END_REF], 87 Sr/ 86 Sr isotope ratios (Roveri et al., 2014b;[START_REF] Gvirtzman | Intra-Messinian truncation surface in the Levant Basin explained by subaqueous dissolution[END_REF][START_REF] Manzi | The onset of the Messinian salinity crisis in the deep Eastern Mediterranean basin[END_REF] and astronomical tuning of the deep seismic record (Ochoa et al., 2015[START_REF] Ochoa | Origin and implications of orbital-induced sedimentary cyclicity in Pliocene well-logs of the Western Mediterranean[END_REF][START_REF] Manzi | The onset of the Messinian salinity crisis in the deep Eastern Mediterranean basin[END_REF][START_REF] Meilijson | Deep-basin evidence resolves a 50-year-old debate and demonstrates synchronous onset of Messinian evaporite deposition in a non-desiccated Mediterranean[END_REF]Meilijson et al., , 2019). Here we focus on the seismic and geological (core-derived) 1 properties of the Upper Unit (and laterally grading/interfingering sediments when present), stratigraphically below the Plio-Quaternary deposits (PQ) suggesting that it belongs to (at least part of) Stage 3.

Western Alborán Basin and westernmost East Alborán Basin

The Alborán Basin has received particular attention because of its proximity to the Gibraltar Corridor [START_REF] Andreetto | Impact of pulsed Atlantic water inflow into the Alboran Basin at the time of the Zanclean flooding[END_REF]Popescu et al., 2015 and references therein). Evaporites only occur on the eastern side of the EAB (which is treated in subsection 4.2; Fig. 2a). To the west of the volcanic archipelago (Booth-Rea et al., 2018, i.e. the WAB) and immediately to the east on the western side of the EAB only terrigenous sediments occur in the MSC interval (Booth-Rea et al., 2018;[START_REF] De La Peña | The evolution of the westernmost Mediterranean basins[END_REF]. Sediments at the Miocene/Pliocene boundary appear in the seismic reflection data as parallel reflectors with increasing reflectivity [START_REF] Comas | Proceedings of the Ocean Drilling Program, Initial Reports[END_REF][START_REF] Comas | The origin and tectonic history of the Alboran Basin: Insights from leg 161 results: Proceedings of the Ocean Drilling Program[END_REF]Booth-Rea et al., 2018). Locally, just below the M discontinuity, some of the reflectors suggest a chaotic seismic facies (Fig. 7a; Booth-Rea et al., 2018;Bulian et al., 2021). Miocene sediments with a maximum thickness of 100 m have been recovered from two out of nine holes drilled in the region (ODP 976B, 978A; [START_REF] Comas | Proceedings of the Ocean Drilling Program, Initial Reports[END_REF][START_REF] Comas | The origin and tectonic history of the Alboran Basin: Insights from leg 161 results: Proceedings of the Ocean Drilling Program[END_REF]. These sediments mostly consist of claystones, siltstones and sandstones with Chondrites and Zoophycos ichnofacies at site 976B and include a conglomerate close to the Miocene/Pliocene boundary at Site 978A. The lack of age-diagnostic fossils hampers their correlation with the Geologic Time Scale (GTS). However, the presence, 1 Lithostratigraphic and biostratigraphic information from DSDP and ODP cores are primarily extracted from the Scientific Shipboard Party documents, accessible from https://www.marum.de/en/Research/Cores-at-BCR.html. These documents are referenced in the text as follows: Ryan et al. (1973): DSDP 120-134; [START_REF] Hsü | Initial reports of the deep sea drilling project: DSDP volume XLII Part[END_REF]: DSDP 371-378; [START_REF] Kastens | Proc. ODP, Init. Repts[END_REF]: ODP 650-656; [START_REF] Comas | Proceedings of the Ocean Drilling Program, Initial Reports[END_REF]: ODP 974-979; [START_REF] Emeis | Proc. ODP, Initial Reports[END_REF]: ODP 963-973.

F. Andreetto et al. high in the Miocene section, of an oligotypic association of ostracods (Candona sp., Loxoconcha muelleri, and Cyprideis sp.) with different stages of growth (Site 978A; [START_REF] Iaccarino | Paleoenvironment of uppermost Messinian sequences in the western Mediterranean (Sites 974, 975, and 978[END_REF] and Paratethyan dinocysts (including Galeacysta etrusca; see subsection 5.2; [START_REF] Popescu | Lago Mare episodes around the Messinian-Zanclean boundary in the deep southwestern Mediterranean[END_REF] indicates a latest Messinian (substage 3.2) age and brackish paleodepositional conditions. Foraminifera and nannofossils are also present, but all species recognized are of no help in narrowing down the paleoenvironmental interpretation because they are considered either definitely or likely to be reworked [START_REF] Iaccarino | Paleoenvironment of uppermost Messinian sequences in the western Mediterranean (Sites 974, 975, and 978[END_REF]. By contrast, [START_REF] Popescu | Lago Mare episodes around the Messinian-Zanclean boundary in the deep southwestern Mediterranean[END_REF] interpreted some species of calcareous nannofossils (Reticulofenestra pseudoumbilicus, Discoaster quinqueramus, Ceratolithus acutus, Triquetrorhabdulus rugosus, Amaurolithus primus) and marine dinocysts as autochthonous.

The nature of the Miocene/Pliocene boundary is also uncertain. According to some authors, the "M" discontinuity is a high-amplitude reflector with evidence of erosion attributed to subaerial processes [START_REF] Andreetto | Impact of pulsed Atlantic water inflow into the Alboran Basin at the time of the Zanclean flooding[END_REF][START_REF] Urgeles | New constraints on the Messinian sealevel drawdown from 3D seismic data of the Ebro margin, western Mediterranean[END_REF] and locally (e.g. close to Site 121; Ryan et al., 1973) associated with an angular unconformity that et al., 2016). Halokinesis of MU gives rise to domes that deform the UU and PQ units. (d) Interpreted seismic profile from the lowermiddle slope of the west Sardinian margin (modified from Dal [START_REF] Dal Cin | Seismic imaging of late Miocene (Messinian) evaporites from western Mediterranean back-arc basins[END_REF]. Thin MU and UU are present on the lower slope, while on the middle slope (and upper slope here not shown) they converge in the margin erosion surface MES. (e) Line drawing of seismic line imaging from the Catalan margin (or Ebro Basin) to the abyss of the Liguro-Provençal basin (modified from Maillard et al., 2011b). Note the pinch out of the MU in the Valencia Basin and of the UU in the Ebro Basin, which is MSC free. (f) Interpreted seismic profile Simbad crossing the depocenter of the CMD showing all the MSC units and erosional surfaces (modified from Raad et al., 2021). (g) Interpreted seismic profile MYS40 illustrating the MU-UU-PQ units in the East-Sardinia Basin and Cornaglia Terrace, separated by the MSC deposits-free Quirra Seamounts (modified from [START_REF] Lymer | The Western Tyrrhenian Sea revisited: New evidence for a rifted basin during the Messinian Salinity Crisis[END_REF].

F. Andreetto et al. abruptly truncates the upper Miocene deposits and morphological highs [START_REF] Comas | The origin and tectonic history of the Alboran Basin: Insights from leg 161 results: Proceedings of the Ocean Drilling Program[END_REF][START_REF] Andreetto | Impact of pulsed Atlantic water inflow into the Alboran Basin at the time of the Zanclean flooding[END_REF][START_REF] Garcia-Castellanos | The Zanclean megaflood of the Mediterranean-Searching for independent evidence[END_REF]. Although the M-reflector was drilled at Sites 976B, 977A and 978A, a lithological contact was only recovered at Site 976B coinciding with a major erosional surface between the early Messinian and the base of the Pliocene (Bulian et al., 2021). Only at Site 978A (and possibly 977A) was a few meters of what may be the contact interval recovered [START_REF] Comas | Proceedings of the Ocean Drilling Program, Initial Reports[END_REF]. This comprises a 25 m-thick cemented succession containing pebbles of volcanic and sedimentary rocks likely to derive from the Alborán substrate (46R, 620.9-630.67 mbsf, between the Pliocene-bearing core 45R and the Messinian-bearing core 47R; [START_REF] Comas | Proceedings of the Ocean Drilling Program, Initial Reports[END_REF]. In contrast, Booth-Rea et al. (2018) concluded that the Mreflector is an unconformity only close to the mud diapirs and owes its erosive shape and angular discordance to the activity of these structures (Fig. 7a). In more undisturbed sectors these authors argue that the boundary is a paraconformity with no evidence of erosion (Fig. 7a). The lack of Messinian erosion in the shallow regions of the WAB margins has prompted the hypothesis that this area did not desiccate during the MSC (Booth-Rea et al., 2018;[START_REF] De La Peña | The evolution of the westernmost Mediterranean basins[END_REF]. This contradicts much of the interpretation made of the DSDP and ODP cores of this interval in the Alborán Sea. The succession recovered by drilling from beneath the Pliocene comprises gravels that contain a mixed Miocene fauna. These sediments and their faunal content are thought to have been reworked from older sediments exposed as Alborán substrate with no evidence of an extensive wet Lago Mare interval immediately before the Zanclean [START_REF] Comas | Proceedings of the Ocean Drilling Program, Initial Reports[END_REF].

Two W-E-aligned erosional channels straddling the Strait of Gibraltar and stretching for 390 km from the easternmost Gulf of Cádiz (Atlantic Ocean) into the Alborán Basin are clearly observed in seismic profiles [START_REF] Garcia-Castellanos | Catastrophic flood of the Mediterranean after the Messinian salinity crisis[END_REF][START_REF] Andreetto | Impact of pulsed Atlantic water inflow into the Alboran Basin at the time of the Zanclean flooding[END_REF]. There is disagreement, however, concerning the timing and nature of their formation. These incisions are classically considered to occur at the very top of the MSC suite (when present) and to be the consequence of the Zanclean megaflood [START_REF] Garcia-Castellanos | Catastrophic flood of the Mediterranean after the Messinian salinity crisis[END_REF], 2020;Estrada et al., 2011 among others). More recently, Krijgsman et al. (2018) highlighted that an accurate age determination for these channel incisions is lacking and that they might have been formed earlier during the MSC as a result of the Mediterranean-Atlantic gateway currents. Interpretation of E-W seismic profiles across the Alborán Basin combined with mammal records in Africa and Iberia led Booth-Rea et al. (2018) to suggest the existence of an emergent volcanic archipelago that temporarily connected southeastern Iberia with northern Africa, separating an open marine, Atlantic-influenced West Alborán Basin realm from a restricted, hydrologically complex Mediterranean realm to the east.

Eastern Alborán, Algero-Balearic and Liguro-Provençal basins

From the eastern margin of the EAB as far east as the Tyrrhenian coast of Italy, the Messinian (evaporites-bearing) trilogy LU-MU-UU is found and sealed by the PQ. The MU and UU are interpreted to fill the deepest depocenters (Algero-Balearic, Valencia and Liguro-Provençal basins; with minor interruptions due to seamounts) and the lower slope domain, where they comprise ~500 to ~800 m of UU and ~1000 m of MU/halite (Figs. 7b-d;[START_REF] Camerlenghi | Morphogenesis of the SW Balearic continental slope and adjacent abyssal plain, Western Mediterranean Sea[END_REF]Lofi et al., 2011aLofi et al., , 2011b;;[START_REF] Geletti | The Messinian Salinity Crisis: new seismic evidence in the West-Sardinian Margin and Eastern Sardo-Provençal basin (West Mediterranean Sea)[END_REF][START_REF] Mocnik | The Messinian Salinity Crisis in the West-Mediterranean Basins: comparison between two rifted margins[END_REF][START_REF] Dal Cin | Seismic imaging of late Miocene (Messinian) evaporites from western Mediterranean back-arc basins[END_REF]Lofi, 2018). Upslope, a thinner, possibly incomplete UU is locally described on the middle-upper continental slopes of Western Corsica [START_REF] Guennoc | West-Corsica Margin: MSC basinal units[END_REF] and Sardinia [START_REF] Mocnik | The Messinian Salinity Crisis in the West-Mediterranean Basins: comparison between two rifted margins[END_REF][START_REF] Dal Cin | Seismic imaging of late Miocene (Messinian) evaporites from western Mediterranean back-arc basins[END_REF] and the northern (Maillard et al., 2006) and southern [START_REF] Maillard | Structure and present-day compression in the offshore area between Alicante and Ibiza Island (Eastern Iberian Margin)[END_REF][START_REF] Mocnik | The Messinian Salinity Crisis in the West-Mediterranean Basins: comparison between two rifted margins[END_REF][START_REF] Dal Cin | Seismic imaging of late Miocene (Messinian) evaporites from western Mediterranean back-arc basins[END_REF] flanks of the Balearic Promontory, even though the structural settings of these locations are mostly dominated by erosion (Fig. 7d). MSC evaporites are absent on the continental shelves bordering the deep Algero-Balearic and Liguro-Provençal basins, where the PQ directly overlies the MES which, in turn, cuts through the middle Miocene deposits [START_REF] Gorini | The Late Messinian salinity crisis and Late Miocene Tectonism: interaction and consequence on the physiography and post-rift evolution of the gulf of Lions margin[END_REF][START_REF] Lofi | Erosional processes and paleo-environmental changes in the Western Gulf of Lions (SW France) during the Messinian Salinity Crisis[END_REF]. The only late Messinian sediments are present as Complex Units (Gulf of Lion, [START_REF] Bessis | Some remarks on the study of subsidence of sedimentary basins. Application to the Gulf of Lions margin (Western Mediterranean)[END_REF][START_REF] Gorini | The Late Messinian salinity crisis and Late Miocene Tectonism: interaction and consequence on the physiography and post-rift evolution of the gulf of Lions margin[END_REF][START_REF] Lofi | Erosional processes and paleo-environmental changes in the Western Gulf of Lions (SW France) during the Messinian Salinity Crisis[END_REF]Algerian Basin, Medaouri et al., 2014;[START_REF] Arab | Tectonostratigraphic evolution of the eastern Algerian margin and basin from seismic data and onshoreoffshore correlation[END_REF]Fig. 7e). CUs can have various origin (Lofi et al., 2011a(Lofi et al., , 2011b)), but when identified at the outlet of drainage systems, they are commonly interpreted as Messinian clastics supplied by rivers [START_REF] Lofi | Erosional processes and paleo-environmental changes in the Western Gulf of Lions (SW France) during the Messinian Salinity Crisis[END_REF]. In the Gulf of Lion, the MES is a high angle unconformity with substantial erosion along highly rugged relief thought to have been generated by fluvial incision [START_REF] Lofi | Erosional processes and paleo-environmental changes in the Western Gulf of Lions (SW France) during the Messinian Salinity Crisis[END_REF]. In contrast, Roveri et al. (2014c) suggested that the drainage networks visible on the seismic could be of subaqueous origin.

When not involved in MU-related deformation processes, the UU appears as a highly reflective series of flat reflectors alternating with less reflective, but concordant, reflectors (Figs. 7b-c;Lofi et al., 2011aLofi et al., , 2011b) ) aggrading in the basin center and onlapping the margins (Fig. 7b; [START_REF] Camerlenghi | Morphogenesis of the SW Balearic continental slope and adjacent abyssal plain, Western Mediterranean Sea[END_REF]Lofi et al., 2011aLofi et al., , 2011b;;[START_REF] Geletti | The Messinian Salinity Crisis: new seismic evidence in the West-Sardinian Margin and Eastern Sardo-Provençal basin (West Mediterranean Sea)[END_REF][START_REF] Mocnik | The Messinian Salinity Crisis in the West-Mediterranean Basins: comparison between two rifted margins[END_REF][START_REF] Dal Cin | Seismic imaging of late Miocene (Messinian) evaporites from western Mediterranean back-arc basins[END_REF]. The aggrading nature, shelf-ward thinning and the onlap terminations of the UU are interpreted as evidence of sedimentation in a lake with fluctuating base level (e.g. [START_REF] Lofi | Erosional processes and paleo-environmental changes in the Western Gulf of Lions (SW France) during the Messinian Salinity Crisis[END_REF]Lofi et al., 2011a). In the abyssal plains (Figs. 2a,7c), nine to ten cycles have been interpreted on high resolution seismic profiles as corresponding to gypsum-marl alternations [START_REF] Geletti | The Messinian Salinity Crisis: new seismic evidence in the West-Sardinian Margin and Eastern Sardo-Provençal basin (West Mediterranean Sea)[END_REF][START_REF] Mocnik | The Messinian Salinity Crisis in the West-Mediterranean Basins: comparison between two rifted margins[END_REF]. At Sites 124 and 372, ~40-50 m of the UU have been drilled at the feet of the east Menorca continental rise and the northern Menorca slope, where 3-4 gypsum-marl cycles are recognized (Fig. 2b; Ryan et al., 1973;Hsü et al., 1978a). Primary gypsum facies are widely overprinted by post-depositional diagenetic processes, but the still recognizable laminated and clastic primary textures indicate precipitation at the water-air interface and emplacement by gravity flows, respectively [START_REF] Lugli | The deep record of the Messinian salinity crisis: evidence of a non-desiccated Mediterranean Sea[END_REF]. The marl interbeds are made from stiff to firm dolomitic mud containing substantial quantities of detrital material intercalated with current-bedded sandstones and, at Site 124, with diatomites (Ryan et al., 1973). Cyprideis sp. specimens are reported from some mudstone interbeds at Site 372, while dwarf planktonic foraminifera are present just below the Miocene/Pliocene boundary at Site 124 (Ryan et al., 1973).

The Miocene/Pliocene boundary coincides with the top of the UU where present (labelled TES when erosional and TS when conformable; Lofi et al., 2011aLofi et al., , 2011b)). In the abyssal plain-lower slope domain it appears to be undulating, although this geometry is related to the deformation of the underlying salt (Figs. 7b-c), and it actually corresponds to a sharp surface lacking signs of erosion (Lofi et al., 2011a(Lofi et al., , 2011b;;[START_REF] Geletti | The Messinian Salinity Crisis: new seismic evidence in the West-Sardinian Margin and Eastern Sardo-Provençal basin (West Mediterranean Sea)[END_REF][START_REF] Mocnik | The Messinian Salinity Crisis in the West-Mediterranean Basins: comparison between two rifted margins[END_REF]. By contrast, the UU-PQ boundary commonly appears strongly erosional in the middle-upper slope and shelf domain, where it coincides with the MES (Fig. 7d; [START_REF] Lofi | Erosional processes and paleo-environmental changes in the Western Gulf of Lions (SW France) during the Messinian Salinity Crisis[END_REF]Maillard et al., 2006;[START_REF] Geletti | The Messinian Salinity Crisis: new seismic evidence in the West-Sardinian Margin and Eastern Sardo-Provençal basin (West Mediterranean Sea)[END_REF][START_REF] Mocnik | The Messinian Salinity Crisis in the West-Mediterranean Basins: comparison between two rifted margins[END_REF]. Among the six DSDP-ODP Sites drilled in this region (Fig. 2b), only Hole 975B recovered the Miocene/Pliocene boundary [START_REF] Iaccarino | Paleoenvironment of uppermost Messinian sequences in the western Mediterranean (Sites 974, 975, and 978[END_REF][START_REF] Marsaglia | Petrography and mineralogy of the uppermost Messinian section and the Pliocene/Miocene boundary at Site 975, Western Mediterranean Sea[END_REF]. Here the Messinian is a few centimeters thick and consists of banded micritic silty clays with minor calcareous siltstones to sandstones typified by a diverse faunal assemblage consisting of dwarf planktonic foraminifera, Ammonia tepida tests and brackish Paratethyan ostracods (Loxocorniculina djafarovi, Euxinocythere praebaquana, Amnicythere idonea, Leptocythere limbate, Candona sp., and Cyprideis sp.; [START_REF] Iaccarino | Paleoenvironment of uppermost Messinian sequences in the western Mediterranean (Sites 974, 975, and 978[END_REF].

Halite is present at the bottom of Hole 134 drilled within the UU (Ryan et al., 1973;[START_REF] Sage | A record of the Messinian salinity crisis on the western Sardinia margin, northwestern Mediterranean[END_REF][START_REF] Lugli | The deep record of the Messinian salinity crisis: evidence of a non-desiccated Mediterranean Sea[END_REF]. High-resolution seismic profiles from both the Algero-Balearic and Ligurian-Provençal basins confirm the presence of a halite layer high in the UU sequence [START_REF] Geletti | The Messinian Salinity Crisis: new seismic evidence in the West-Sardinian Margin and Eastern Sardo-Provençal basin (West Mediterranean Sea)[END_REF][START_REF] Mocnik | The Messinian Salinity Crisis in the West-Mediterranean Basins: comparison between two rifted margins[END_REF]. This layer is up to 50 m thick [START_REF] Dal Cin | Seismic imaging of late Miocene (Messinian) evaporites from western Mediterranean back-arc basins[END_REF] and is correlated with an erosional surface (called IES: Intermediate Erosional Surface by Lofi et al., 2011aLofi et al., , 2011b) associated with an angular unconformity which is better developed on the lower slope (Fig. 7d). [START_REF] Geletti | The Messinian Salinity Crisis: new seismic evidence in the West-Sardinian Margin and Eastern Sardo-Provençal basin (West Mediterranean Sea)[END_REF] and [START_REF] Mocnik | The Messinian Salinity Crisis in the West-Mediterranean Basins: comparison between two rifted margins[END_REF] interpreted this layer as autochtonous and indicative of at least one important sea level drop during UU deposition. However, this intra UU halite layer is always described from areas strongly affected by salt F. Andreetto et al. diapirism (just like in the Ionian Abyssal Plain; see subsection 4.6.1) and is never found in adjacent, undisturbed areas (see [START_REF] Camerlenghi | Morphogenesis of the SW Balearic continental slope and adjacent abyssal plain, Western Mediterranean Sea[END_REF][START_REF] Geletti | The Messinian Salinity Crisis: new seismic evidence in the West-Sardinian Margin and Eastern Sardo-Provençal basin (West Mediterranean Sea)[END_REF][START_REF] Mocnik | The Messinian Salinity Crisis in the West-Mediterranean Basins: comparison between two rifted margins[END_REF][START_REF] Dal Cin | Seismic imaging of late Miocene (Messinian) evaporites from western Mediterranean back-arc basins[END_REF], two features that may suggest it has an allochthonous origin.

Site 134 shows evidence of a "desiccation crack" cutting through a sandy silt layer interbedded with unaffected laminated halite (Hsü et al., 1973c). Unfortunately, the core photograph of this crack has been published in two different orientations (Hsü et al., 1973a(Hsü et al., , 1973b)), leading both Hardie and Lowenstein (2004) and [START_REF] Lugli | The deep record of the Messinian salinity crisis: evidence of a non-desiccated Mediterranean Sea[END_REF] to question the evidence for subaerial desiccation. Because of these ambiguities, we suggest to dismiss this example from the book of evidence.

Valencia Basin

The Valencia Basin (VB; Fig. 2a) is an aborted rift formed during the late Oligocene-early Miocene opening of the back-arc Liguro-Provençal Basin (e.g. [START_REF] Jolivet | Lithospheric-scale geodynamic context of the Messinian Salinity Crisis[END_REF]. It is located between the Spanish Ebro Margin to the west and the Balearic Promontory to the east, while it grades into the Liguro-Provençal Basin to the E/NE (Fig. 7e; Maillard and Mauffret, 2006;Maillard et al., 2006).

Numerous exploratory boreholes exist on the western Ebro margin. These boreholes, tied to seismic data, confirm that MSC-related sediments on the northwestern shelf are missing (Fig. 7e; [START_REF] Urgeles | New constraints on the Messinian sealevel drawdown from 3D seismic data of the Ebro margin, western Mediterranean[END_REF][START_REF] Pellen | The Messinian Ebro River incision[END_REF]. The only MSC feature present is a prominent erosional surface (the MES) incising Serravallian-early Messinian sediments [START_REF] Urgeles | New constraints on the Messinian sealevel drawdown from 3D seismic data of the Ebro margin, western Mediterranean[END_REF]. By contrast, on the southwestern and southern part of the margin, well data show the presence of evaporitic sediments (e.g. Delta L and Golfo de Valencia D1 boreholes;[START_REF] Del Olmo | The Messinian in the Gulf of Valencia and Alboran Sea (Spain): paleogeography and paleoceanography implications[END_REF][START_REF] Del Olmo | The Messinian record of Spanish onshore and offshore data (Atlantic Ocean and Western Mediterranean Sea)[END_REF][START_REF] Lozano | Astrobiochronological Constraints on Margin to deep basin correlations across the Balearic Promontory and the Valencia basin[END_REF]. Del Olmo and Martín (2016) described these evaporites as primary selenites and ascribed them to MSC Stage 1 (their unit M7). [START_REF] Lozano | Astrobiochronological Constraints on Margin to deep basin correlations across the Balearic Promontory and the Valencia basin[END_REF] described the same evaporitic deposits in the same boreholes as 'white anhydrites', leaving open the question as to whether the anhydrite is primary (sabhka's) or secondary at the expense of a primary gypsum facies. On the eastern margin of the VB boreholes and seismic studies suggest there are no MSC units with only a prominent MES cutting pre-MSC sediments (Driussi et al., 2015;Raad et al., 2021). All authors conclude that the shelves of VB were exposed to subaerial erosion during and following the main drawdown.

MSC deposits are also absent along the slopes and, where present, consist of coarse-and fine-grained terrigenous facies filling valleys largely related to fluvial incision (Fig. 7e; [START_REF] Stampfli | Messinian paleorelief from a 3D seismic survey in the Tarrasco concession area (Spanish Mediterranean Sea)[END_REF].

A different situation features in the depocenter. Despite its presentday depth of > 2000 m, no MU is observed in the depocenter, as the salt pinches-out at the edge of the deeper Provençal Basin (Fig. 7e). Only the seismic properties of UU suggest that it is roughly continuous from the Provençal Basin into the VB (Fig. 7e; see subsection 4.2), although it thins from ~1000 m to < 500 m. The UU is characterized by aggrading and onlapping geometries towards the slopes, where it also thins out until it disappears along the middle-upper slope (Fig. 7e; Maillard et al., 2006;[START_REF] Cameselle | Large-scale margin collapse during Messinian early sea-level drawdown: the SW Valencia trough, NW Mediterranean[END_REF]. Maillard et al. (2006), [START_REF] Urgeles | New constraints on the Messinian sealevel drawdown from 3D seismic data of the Ebro margin, western Mediterranean[END_REF][START_REF] Cameselle | Late Miocene sedimentary architecture of the Ebro Continental Margin (Western Mediterranean): implications to the Messinian Salinity Crisis[END_REF] and [START_REF] Cameselle | Large-scale margin collapse during Messinian early sea-level drawdown: the SW Valencia trough, NW Mediterranean[END_REF] all stated that the UU formed during an important Mediterranean-level lowstand (~1000 m). Several Complex Units (CU), with different origin, have been observed and described as belonging to the MSC [START_REF] Cameselle | Large-scale margin collapse during Messinian early sea-level drawdown: the SW Valencia trough, NW Mediterranean[END_REF].

DSDP Site 122, drilled at the mouth of a valley incision, recovered a few meters of sand-gravels made of well-rounded basalt, marine limestones, nodules of crystalline gypsum and mollusk fragments in a claysilty matrix rich in deep-water benthonic foraminifera and early Pliocene nannofossils, all interpreted as erosional products of the VB seabed (Ryan et al., 1973). The uppermost Messinian in two industrial wells (Ibiza Marino and Cabriel boreholes; see [START_REF] Lozano | Astrobiochronological Constraints on Margin to deep basin correlations across the Balearic Promontory and the Valencia basin[END_REF] is represented by intercalations of clastic gypsum/anhydrite and marls (unit M8-P1 of Del [START_REF] Del Olmo | The Messinian record of Spanish onshore and offshore data (Atlantic Ocean and Western Mediterranean Sea)[END_REF]. These are interpreted as turbidites sourced from the shelf and/or slope during a lowstand phase of the Mediterranean base level [START_REF] Del Olmo | The Messinian in the Gulf of Valencia and Alboran Sea (Spain): paleogeography and paleoceanography implications[END_REF][START_REF] Del Olmo | The Messinian record of Spanish onshore and offshore data (Atlantic Ocean and Western Mediterranean Sea)[END_REF][START_REF] Cameselle | Large-scale margin collapse during Messinian early sea-level drawdown: the SW Valencia trough, NW Mediterranean[END_REF].

In seismic profiles, the UU/PQ transition (M-reflector or TES) is locally both sharp and smooth (in more distal settings) and erosional (in more proximal settings; Fig. 7e). Maillard and Mauffret (2006) indicate that the smooth parts have been caused by increasing fresh water influx during the Lago-Mare phase, leading to dissolution of the evaporites, and the rough erosional segments are of subaerial origin. For [START_REF] Cameselle | Large-scale margin collapse during Messinian early sea-level drawdown: the SW Valencia trough, NW Mediterranean[END_REF], the top of the UU is a smooth and conformable downlap surface, representing the rapid inundation of the basin with only local minor erosional features.

Balearic Promontory

Sticking out from the surrounding deep-water locations, the Balearic Promontory (BP; Fig. 2a) is a continental high that has undergone tectonic extension since the late Serravallian [START_REF] Roca | The Neogene structure of the eastern iberian margin: structural constraints on the crustal evolution of the Valencia trough (western mediterranean)[END_REF][START_REF] Sabat | Geological structure and evolution of Majorca: implications for the origin of the Western Mediterranean[END_REF]. During the Messinian, it comprised in topographic lows/subbasins at different water depths and separated by structural highs/sills [START_REF] Maillard | Record of the Messinian Salinity Crisis in the SW Mallorca area (Balearic Promontory, Spain)[END_REF]Driussi et al., 2015;Roveri et al., 2019b;Raad et al., 2021). The area has been the subject of multiple studies [START_REF] Maillard | Record of the Messinian Salinity Crisis in the SW Mallorca area (Balearic Promontory, Spain)[END_REF]Driussi et al., 2015;Ochoa et al., 2015;Roveri et al., 2019b;Raad et al., 2021) and several controversies arose after the publication of Roveri et al. (2019b).

The first controversy concerns the Messinian paleodepth of the BP's subbasins. According to Roveri et al. (2019b) the subbasins were shallow during the Messinian and acquired today's paleodepths following a strong post-MSC subsidence; [START_REF] Maillard | Atlas of the Messinian Salinity Crisis markers in the Mediterranean and Black SeasMémoires de la Société géologique de France 179[END_REF], [START_REF] Maillard | Record of the Messinian Salinity Crisis in the SW Mallorca area (Balearic Promontory, Spain)[END_REF] and Raad et al. (2021), instead, consider the tectonic movements in the BP to have been minor since the late Miocene (Messinian) and the region to have been already structured as it is today during the MSC. Well-to-seismic ties in the shallower basins closer to the Spanish coast (i.e. Bajo Segura, San Pedro and Elche basins) comprise up to 14 Stage 1 primary gypsum-marl cycles similar to the onshore PLG unit (Lugli et al., 2010) truncated at the top by the MES (Soria et al., 2008a(Soria et al., , 2008b;;Ochoa et al., 2015). At first, Ochoa et al. (2015) concluded that all sediments overlying the MES are Pliocene in age. A later re-appraisal of the same downhole logging data and cuttings led [START_REF] Ochoa | Origin and implications of orbital-induced sedimentary cyclicity in Pliocene well-logs of the Western Mediterranean[END_REF] to attribute the first ~13 m-thick micritic and evaporite-free sediments to the late Messinian (stage 2 or 3 of the MSC according to the authors). The MSC stratigraphy of the shallowest offshore basins of the BP closely resembles that described from cores and outcrops onshore Mallorca (see subsection 3.4;Roveri et al., 2019b).

High resolution seismic reflection data in the Central Mallorca Depression (CMD) highlighted up to 500 m of MSC deposits made of a Bedded Unit (BU) and a thin salt layer [START_REF] Maillard | Record of the Messinian Salinity Crisis in the SW Mallorca area (Balearic Promontory, Spain)[END_REF]Driussi et al., 2015). This BU has never been drilled and, therefore, lacks lithological and chronostratigraphic constraints. Two contrasting chronostratigraphic and lithological interpretations are proposed: Roveri et al. (2019b) ascribed these sediments to Stage 2 and 3 and suggested that only reworked evaporites and halite are present. By contrast, following the seismostratigraphic description of [START_REF] Maillard | Record of the Messinian Salinity Crisis in the SW Mallorca area (Balearic Promontory, Spain)[END_REF], Ochoa et al. (2015) and Raad et al. (2021) inferred the presence of Stage 1 gypsum also in the CMD. Raad et al. (2021) made a step forward by disclosing a possible tripartition of the BU unit (Fig. 7f). In their seismostratigraphic framework, Raad et al. (2021) noticed that the uppermost evaporite-bearing unit (called BU3), ~120 m-thick in the CMD, has geometric, stratigraphic and facies analogies with the astronomically-tuned UG unit of the Caltanissetta Basin (Fig. 3a) that endorse its attribution to Stage 3. Similar to the UU in the deepest basins (see subsection 4.2), BU3 consists of up to 9 low-and medium-amplitude reflectors that are interpreted as alternating terrigenous and gypsum beds [START_REF] Maillard | Record of the Messinian Salinity Crisis in the SW Mallorca area (Balearic Promontory, Spain)[END_REF]Raad et al., 2021). Reflectors are parallel and continuous in the more distal areas, while they appear more chaotic in the more proximal sectors (Raad et al., 2021). The base of BU3 coincides with the erosional top of the salt, F. Andreetto et al. interpreted as created by salt exposure, dissolution and locally salt gliding towards the depocenter (Fig. 7f; Raad et al., 2021). By contrast, the top of BU3, which corresponds to the Miocene/Pliocene boundary, is largely flat without signs of erosion (Fig. 7f; [START_REF] Maillard | Record of the Messinian Salinity Crisis in the SW Mallorca area (Balearic Promontory, Spain)[END_REF]Raad et al., 2021). An irregular geometry is sometimes visible, but is likely to be related to deformation of the underlying salt (Fig. 7f; Raad et al., 2021).

Tyrrhenian Basin

The Tyrrhenian Basin to the east of Sardinia is a back-arc basin that opened by continental rifting and oceanic spreading related to the eastward migration of the Apennine subduction system from middle Miocene to Pliocene times [START_REF] Gaullier | Salt tectonics and crustal tectonics along the Eastern Sardinian margin, Western Tyrrhenian: new insights from the[END_REF][START_REF] Lymer | The Western Tyrrhenian Sea revisited: New evidence for a rifted basin during the Messinian Salinity Crisis[END_REF]Loreto et al., 2020 and references therein). Three main domains are traditionally identified [START_REF] Lymer | The Western Tyrrhenian Sea revisited: New evidence for a rifted basin during the Messinian Salinity Crisis[END_REF] and references therein): 1) the East Sardinia Basin, with present-day water depths between 200-2000 m consisting of a system of seamounts and depressions that do not contain MSC sediments [START_REF] Lymer | The Western Tyrrhenian Sea revisited: New evidence for a rifted basin during the Messinian Salinity Crisis[END_REF]; 2) the Cornaglia Terrace (2000-3000 m deep), a wide, flat area with dispersed structural highs; 3) the Tyrrhenian Basin s.s., with water depths varying from 3000-3600 m. Whether the Tyrrhenian Basin acquired the present-day topography and hypsometry before the MSC [START_REF] Lymer | The Western Tyrrhenian Sea revisited: New evidence for a rifted basin during the Messinian Salinity Crisis[END_REF] or at least part of it (e.g. Eastern Sardinia margin, where Site 654 is located, and Northern Tyrrhenian) was much shallower (possibly comparable to the Caltanissetta Basin; Roveri et al., 2014b) and underwent extension and subsidence during the Messinian-Pliocene (e.g. [START_REF] Kastens | The geological evolution of the Tyrrhenian Sea: An introduction to the scientific results of ODP Leg 107[END_REF][START_REF] Loreto | Extensional tectonics during the Tyrrhenian back-arc basin formation and a new morphotectonic map[END_REF]) is still unresolved.

The MSC units in seismic profiles from the Tyrrhenian Basin (Fig. 7g) are very similar to the ones described in the Algero-Balearic and Liguro-Provençal basins (Fig. 7b-c; [START_REF] Gaullier | Salt tectonics and crustal tectonics along the Eastern Sardinian margin, Western Tyrrhenian: new insights from the[END_REF][START_REF] Lymer | The Western Tyrrhenian Sea revisited: New evidence for a rifted basin during the Messinian Salinity Crisis[END_REF]. ODP Sites 652, 653 and 654 confirmed the seismic-inferred lithological nature of UU as consisting, of gypsum-mudstone alternations (8 are counted at Site 654; [START_REF] Kastens | Proc. ODP, Init. Repts[END_REF][START_REF] Borsetti | Messinian and pre-Messinian sediments from ODP Leg 107 Sites 652 and 654 in the Tyrrhenian Sea: sedimentologic and petrographic study and possible comparisons with Italian sequences[END_REF]Roveri et al., 2014b). Intercalations of ripple-cross-laminated, fine-grained, azoic sandstones occur within the mudstone intervals in places [START_REF] Cita | Messinian Paleoenvironments: New Evidence from the Tyrrhenian Sea (ODP LEG 107)[END_REF][START_REF] Iaccarino | Paleoenvironment of uppermost Messinian sequences in the western Mediterranean (Sites 974, 975, and 978[END_REF]. In some mudstone samples, the ostracod Cyprideis sp. (Site 654) and Candona sp. (DSDP Hole 974B) and the foraminifera Ammonia beccarii and Ammonia tepida have been found, indicating a shallow-water (< 50 m) brackish environment (see subsections 5. 1 and 5.4;[START_REF] Cita | Messinian Paleoenvironments: New Evidence from the Tyrrhenian Sea (ODP LEG 107)[END_REF][START_REF] Iaccarino | Paleoenvironment of uppermost Messinian sequences in the western Mediterranean (Sites 974, 975, and 978[END_REF]. 87 Sr/ 86 Sr isotope ratios of UU gypsum and planktic foraminifera of the overlying Pliocene (Unit 1 at Site 654) show values much lower (from 0.708627 to 0.708745) and roughly equivalent (from 0.708935 to 0.709112) to coeval ocean water (~0. 709020-30;McArthur et al., 2012), respectively (Müller et al., 1990;[START_REF] Müller | Origin and age of the Mediterranean Messinian evaporites: implications from Sr isotopes[END_REF]. Similar 87 Sr/ 86 Sr values were obtained from the gypsum cored at Site 652 (0.708626-0.708773;[START_REF] Müller | Origin and age of the Mediterranean Messinian evaporites: implications from Sr isotopes[END_REF].

The Miocene/Pliocene boundary at DSDP Site 132 is placed above a cross-bedded sand rich in quartz, mica, pyrite, rounded fragments of gypsum and specimens of Ammonia beccarii and Elphidium macellus (Ryan et al., 1973). In the adjacent ODP Site 653 a similar sandstone is found slightly below the biostratigraphically-defined Messinian/Pliocene boundary and ~70 cm of grey mudstones with foraminifera and nannofossils of undisclosed provenance are squeezed in between [START_REF] Cita | Messinian Paleoenvironments: New Evidence from the Tyrrhenian Sea (ODP LEG 107)[END_REF]. These mudstones also contain rare dwarf planktic foraminifera (Globorotalia acostaensis, Orbulina universa, and Globigerina bulbosa; [START_REF] Cita | Messinian Paleoenvironments: New Evidence from the Tyrrhenian Sea (ODP LEG 107)[END_REF].

Overall, the uppermost Messinian sediments of the Tyrrhenian Basin are interpreted as having been deposited in lakes with periodic episodes of increased salinity and dilution under the strong influence of high energy rivers and, perhaps ocassionally, of the Atlantic [START_REF] Cita | Messinian Paleoenvironments: New Evidence from the Tyrrhenian Sea (ODP LEG 107)[END_REF]Müller et al., 1990;[START_REF] Müller | Origin and age of the Mediterranean Messinian evaporites: implications from Sr isotopes[END_REF].

Ionian Basin

The deepest basin in the Mediterranean today is the Ionian Basin, with its lowest point at 5,267 meters. The so-called Ionian Abyssal Plain (IAP) is bounded on all sides by pre-MSC structural highs (Fig. 2a; [START_REF] Camerlenghi | Seismic markers of the Messinian salinity crisis in the deep Ionian Basin[END_REF]: the Malta Escarpment to the west; the Medina Escarpment to the south separating it from the Gulf of Sirt (Fig. 8a); the Gargano-Pelagosa and/or Otranto sill to the north dividing it from the Adriatic Foreland and, finally, an unnamed sill to the east, separating the IAP from the Levant Basin. Evidence from recent high-resolution seismic studies across the region have been used to support Stage 3 desiccation models (e.g. Hsü et al., 1978a[START_REF] Hsü | Initial reports of the deep sea drilling project: DSDP volume XLII Part[END_REF][START_REF] Bowman | A comprehensive review of the MSC facies and their origins in the offshore Sirt Basin, Libya[END_REF][START_REF] Micallef | Evidence of the Zanclean megaflood in the eastern Mediterranean Basin[END_REF][START_REF] Micallef | Geomorphic evolution of the Malta Escarpment and implications for the Messinian evaporative drawdown in the eastern Mediterranean Sea[END_REF][START_REF] Camerlenghi | Seismic markers of the Messinian salinity crisis in the deep Ionian Basin[END_REF][START_REF] Spatola | A single-stage megaflood at the termination of the Messinian salinity crisis: geophysical and modelling evidence from the eastern Mediterranean Basin[END_REF].

Ionian Abyssal Plain

The typical "trilogy" of seismic units representing the MSC deposition in the Western Mediterranean is recognized also in the IAP by [START_REF] Gallais | Seismic Atlas of the Messinian salinity crisis markers in the Mediterranean sea. Volume 2[END_REF] and [START_REF] Mocnik | Seismic Atlas of the Messinian salinity crisis markers in the Mediterranean sea. Volume 2[END_REF]. However, [START_REF] Camerlenghi | Seismic markers of the Messinian salinity crisis in the deep Ionian Basin[END_REF] states the MSC sequence in the IAP is ~1,300 m-thick and composed of only two units (Fig. 8a). The lowermost 150-700 m-thick Mobile Unit (MU) is seismically transparent without discernible bedding and with diapiric structures, all features diagnostic of halite. By contrast, the 350-1,000 m-thick Upper Unit (UU) alternates highly reflective with acoustically transparent reflectors (Figs. 8a-b), similar to those described of the UU sequences of the Western Mediterranean (Figs. 7bc). These are therefore assumed to represent gypsum-mudstone cycles [START_REF] Camerlenghi | Seismic markers of the Messinian salinity crisis in the deep Ionian Basin[END_REF]. The uppermost 80 m of UU has been recovered from DSDP Site 374 [START_REF] Hsü | Initial reports of the deep sea drilling project: DSDP volume XLII Part[END_REF], confirming the presence of gypsum (both primary cumulate and clastic; [START_REF] Lugli | The deep record of the Messinian salinity crisis: evidence of a non-desiccated Mediterranean Sea[END_REF] alternating with mudstones (Unit III; [START_REF] Hsü | Initial reports of the deep sea drilling project: DSDP volume XLII Part[END_REF]. These mudstones are largely barren of in situ fossils. However, the presence of some foraminifera and siliceous microfossils led Cita et al. (1978) and Hsü et al. (1978a) to suggest that sporadic marine incursions, possibly from the Indian Ocean, took place during Stage 3. Similar to Site 372, the basal part of Hole 374 intercepted one thin (~3.5 m) halite layer within the UU [START_REF] Hsü | Initial reports of the deep sea drilling project: DSDP volume XLII Part[END_REF].

The UU and the overlying Zanclean (subunit PQc of [START_REF] Camerlenghi | Seismic markers of the Messinian salinity crisis in the deep Ionian Basin[END_REF] reflectors are conformably folded throughout most of the abyssal plain, locally showing chaotic internal structure, irregular folding mimicking V-shaped incisions and truncations (Fig. 8b; [START_REF] Camerlenghi | Seismic markers of the Messinian salinity crisis in the deep Ionian Basin[END_REF]. These features are interpreted by [START_REF] Camerlenghi | Seismic markers of the Messinian salinity crisis in the deep Ionian Basin[END_REF] as fluvial valleys carved in subaerially exposed evaporites by the Eso-Sahabi fluvial system, the closest fluvial drainage network to the area (see [START_REF] Micallef | Evidence of the Zanclean megaflood in the eastern Mediterranean Basin[END_REF] that drained Libya in the late Miocene [START_REF] Griffin | Aridity and humidity: two aspects of the late Miocene climate of North Africa and the Mediterranean[END_REF] and has been traced across the Gulf of Sirt offshore [START_REF] Sabato Ceraldi | Messinian seismic facies in offshore Sirt Basin Libya and implications for sub-Messinian seismic imaging[END_REF][START_REF] Bowman | A comprehensive review of the MSC facies and their origins in the offshore Sirt Basin, Libya[END_REF]. Several arguments counteract this interpretation: 1) the coherent, deformation, mostly of post-Messinian age, of both the UU and the lower Zanclean; 2) the absence of fluvial facies above the bottom of the "valleys", which instead correspond to a mudstone interval that underwent post-depositional dolomitization (Unit II; see below; Fig. 8b); 3) the unlikelihood that the Eso-Sahabi fluvial system managed to bypass the Medina Ridge divide (Fig. 8a). Alternatively, the apparent incisions at the M/P boundary in the IAP may represent post-sedimentary folds resulting from post-Messinian tectonic and/or salt movements-driven deformation (e.g. [START_REF] Mocnik | Seismic Atlas of the Messinian salinity crisis markers in the Mediterranean sea. Volume 2[END_REF]. At Site 374 the Miocene/Pliocene boundary has been recovered (Unit II), but its primary nature (likely a mudstone) is obscured by diagenesis (cementation by dolomite; [START_REF] Hsü | Initial reports of the deep sea drilling project: DSDP volume XLII Part[END_REF]. A lithified dolostone at the (seismic) predicted depth of the M-Horizon is a characteristic of several sites. Sometimes this interval has been recovered (e.g. Sites 125 and 374; Ryan et al., 1973;[START_REF] Hsü | Initial reports of the deep sea drilling project: DSDP volume XLII Part[END_REF][START_REF] Comas | Proceedings of the Ocean Drilling Program, Initial Reports[END_REF]; at others the hard lithology is inferred by the torqueing of the drill string (resistance to turning) accompanied by bouncing of the drill bit at the (e.g. Sites 122, 124, 125, 132, 133 and 134;Ryan et al., 1973). Dolomitization was (Hsü et al., 1973a(Hsü et al., , 1973b;;Ryan et al., 1973) and still is (e.g. Ryan, 2009) largely considered a "diagnostic feature of tidal (sabkha) sediments" (Friedman, 1973, pp. 705). Its occurrence at locations with present-day water depth exceeding 2000 m was therefore considered strong evidence that the Mediterranean floor was subaerially exposed prior to the Zanclean marine replenishment (e.g. Ryan et al., 1973). It is, however, now widely accepted that dolomite precipitation is not exclusive of sabkha environments, but rather is a common process also in bottom sediments under a relatively deep water column (see Dela Pierre et al., 2012, 2014 andreferences therein). In the specific case of the offshore Mediterranean's M/P boundary on the Ionian Abyssal Plain, already in the '70s dolomitization was thought to have occurred after burial [START_REF] Hsü | Initial reports of the deep sea drilling project: DSDP volume XLII Part[END_REF], a conclusion recently reinforced by [START_REF] Mckenzie | Subsurface dolomite formation during post-depositional flow of sulphate-bearing fluids from underlying salt giants: Early Pliocene example at DSDP Leg 42A, Site 374, Ionian Abyssal Plain[END_REF].

Malta Escarpment

At the foot of the Malta Escarpment, [START_REF] Micallef | Evidence of the Zanclean megaflood in the eastern Mediterranean Basin[END_REF][START_REF] Micallef | Geomorphic evolution of the Malta Escarpment and implications for the Messinian evaporative drawdown in the eastern Mediterranean Sea[END_REF] and [START_REF] Spatola | A single-stage megaflood at the termination of the Messinian salinity crisis: geophysical and modelling evidence from the eastern Mediterranean Basin[END_REF] amalgamated the MU and UU into one seismic unit, Unit 3, which is separated from the Plio-Quaternary marine sediments (Unit 1) by Unit 2, a chaotic transparent seismic package (Fig. 8c). Unit 2 has a maximum thickness of 760-860 m, a volume of 1430-1620 km 3 and a wedge-shaped geometry that thins eastwards, disappearing before reaching the IAP [START_REF] Micallef | Evidence of the Zanclean megaflood in the eastern Mediterranean Basin[END_REF]. [START_REF] Micallef | Evidence of the Zanclean megaflood in the eastern Mediterranean Basin[END_REF] and [START_REF] Spatola | A single-stage megaflood at the termination of the Messinian salinity crisis: geophysical and modelling evidence from the eastern Mediterranean Basin[END_REF] proposed a lithological/sedimentological interpretation of this chaotic body, suggesting it is composed of well-sorted sediments of the Pelagian Platform to the west of the Malta Escarpment, with coarser material at the mouth grading into more distal finer sediments. This chaotic body has recently been attributed to the Zanclean megaflood during its passage from the western to the eastern Mediterranean via a gateway located in south-eastern Sicily [START_REF] Micallef | Evidence of the Zanclean megaflood in the eastern Mediterranean Basin[END_REF][START_REF] Micallef | Geomorphic evolution of the Malta Escarpment and implications for the Messinian evaporative drawdown in the eastern Mediterranean Sea[END_REF][START_REF] Spatola | A single-stage megaflood at the termination of the Messinian salinity crisis: geophysical and modelling evidence from the eastern Mediterranean Basin[END_REF]. Given the rapidity of the reflooding (≤ 2 years, [START_REF] Garcia-Castellanos | Catastrophic flood of the Mediterranean after the Messinian salinity crisis[END_REF], 2020), this interpretation implies rapid mass deposition. Other interpretations of this Unit 2 include being the result of extensive marine mass movement [START_REF] Polonia | The Calabrian Arc subduction complex in the Ionian Sea: Regional architecture, active deformation, and seismic hazard[END_REF], being folded UU [START_REF] Butler | Stratigraphic variations control deformation patterns in evaporite basins: Messinian examples, onshore and offshore Sicily (Italy)[END_REF] or being a complex unit built during lower sea level phases (Lofi et al., 2011a(Lofi et al., , 2011b)).

Gulf of Sirt

The Gulf of Sirt (or Sirt embayment; Figs. 2a, 8a), the offshore extension of the Sirt Basin onshore Libya (Griffin, 2002 and references therein), is cross-cut by high-density seismic and well datasets as a result of oil potential of the region [START_REF] Fiduk | Evaporites, petroleum exploration, and the Cenozoic evolution of the Libyan shelf margin, central North Africa[END_REF]. However, only few studies have tackled the MSC (e.g. [START_REF] Hallett | Petroleum Geology of Libya[END_REF][START_REF] Fiduk | Evaporites, petroleum exploration, and the Cenozoic evolution of the Libyan shelf margin, central North Africa[END_REF][START_REF] Bowman | A comprehensive review of the MSC facies and their origins in the offshore Sirt Basin, Libya[END_REF].

In the Sirt embayment the MSC unit(s) is unevenly distributed in subbasins controlled by a pre-existing topography, there is little distinction between the MU and UU, the overall thickness of the MSC unit(s) is lower and the degree of deformation is higher than in the adjacent IAP (Fig. 8a; [START_REF] Camerlenghi | Seismic markers of the Messinian salinity crisis in the deep Ionian Basin[END_REF]. The presence of halite in the Sirt embayment is debated, but most authors think it is absent (see [START_REF] Fiduk | Evaporites, petroleum exploration, and the Cenozoic evolution of the Libyan shelf margin, central North Africa[END_REF][START_REF] Sabato Ceraldi | Messinian seismic facies in offshore Sirt Basin Libya and implications for sub-Messinian seismic imaging[END_REF]Lofi, 2018;[START_REF] Jagger | Tectono-stratigraphic evolution of the SE Mediterranean passive margin, offshore Egypt and Libya[END_REF]Fig. 2b). [START_REF] Bowman | A comprehensive review of the MSC facies and their origins in the offshore Sirt Basin, Libya[END_REF] distinguishes seven seismic units within the MSC-related sequence (Fig. 8d). On the basis of high-resolution 3D and 2D data, each seismic unit has been interpreted consisting of clastics filling erosional channels cutting up to 100 m deep and wide (Fig. 8d) and evaporites (gypsum and anhydrite) alternating with precessional cyclicity [START_REF] Bowman | A comprehensive review of the MSC facies and their origins in the offshore Sirt Basin, Libya[END_REF]. The presence of anhydrite in the topmost part of the sequence is confirmed by the B1 NC 35A well [START_REF] Hallett | Petroleum Geology of Libya[END_REF]. Sabato [START_REF] Sabato Ceraldi | Messinian seismic facies in offshore Sirt Basin Libya and implications for sub-Messinian seismic imaging[END_REF] and [START_REF] Bowman | A comprehensive review of the MSC facies and their origins in the offshore Sirt Basin, Libya[END_REF] envisaged a three-step evolution of each unit: 1) evaporitic deposition during precession maxima in a dried out Sirt embayment; 2) erosion of the valleys during the arid-wet transition fed by the Eso-Sahabi paleofluvial system; 3) filling of the valleys with the fluvial sediments during the wet phase. Based on this chronostratigraphic interpretation, the evaporite cycles should be time equivalent to most of Stage 3, with the upper four seismic units reflecting the Lago-Mare phase (Fig. 1a).

Levant Basin

Abyssal plain

Old seismic data in the Levant Basin show an up to 2 km-thick, high velocity, acoustically transparent sequence bounded by the N-reflector at the base and the M-reflector at the top (Figs. 8e-f;[START_REF] Ryan | Messinian badlands on the southeastern margin of the Mediterranean Sea[END_REF][START_REF] Netzeband | The structural evolution of the Messinian evaporites in the Levantine Basin[END_REF]. This sequence thickens and extends for tens of kilometers towards WNW and thins eastward along the continental margin (Fig. 8e), where the N and M-reflectors converge forming the condensed MSC section of the Mavqiim and Afiq formations (described in subsection 4.7.2; [START_REF] Gvirtzman | Intra-Messinian truncation surface in the Levant Basin explained by subaqueous dissolution[END_REF][START_REF] Manzi | The onset of the Messinian salinity crisis in the deep Eastern Mediterranean basin[END_REF]. High resolution 2D and 3D industrial seismic and well data from the southern Levant Basin revealed that this transparent sequence is largely made of pure halite with internal stratification picked out by diatomite, clay-and clastic-rich layers that allowed the division of the sequence into six subunits, basinward-tilted and truncated at the top by the flat TES (Fig. 8f; [START_REF] Gvirtzman | Intense salt deformation in the Levant Basin in the middle of the Messinian salinity crisis[END_REF][START_REF] Gvirtzman | Bathymetry of the Levant basin: interaction of salt-tectonics and surficial mass movements[END_REF][START_REF] Gvirtzman | Intra-Messinian truncation surface in the Levant Basin explained by subaqueous dissolution[END_REF][START_REF] Feng | Lithology and characteristics of the Messinian evaporite sequence of the deep Levant Basin, Eastern Mediterranean[END_REF][START_REF] Feng | Intra-salt deformation: implications for the evolution of the Messinian evaporites in the levant basin, eastern Mediterranean[END_REF][START_REF] Manzi | The onset of the Messinian salinity crisis in the deep Eastern Mediterranean basin[END_REF][START_REF] Meilijson | Deep-basin evidence resolves a 50-year-old debate and demonstrates synchronous onset of Messinian evaporite deposition in a non-desiccated Mediterranean[END_REF]Meilijson et al., , 2019)). Clastic beds ~3 to 20 m-thick are abundant in the highly reflective and chaotic Unit 5 (i.e. Interbedded evaporites of Meilijson et al., 2019;MC2 unit of Feng et al., 2016;Figs. 3b, 4a), where they are interbedded with evaporites (probably halite with minor occurrences of anhydrite) varying in thickness from ~6 to 30 m [START_REF] Manzi | The onset of the Messinian salinity crisis in the deep Eastern Mediterranean basin[END_REF]Meilijson et al., 2019). Paleontological analyses of cuttings probably belonging to one of the clastic beds revealed the presence of a few mollusk fragments, ostracods, echinoid spines and a relatively rich assemblage of benthic and planktic foraminifera which Meilijson et al. (2019) concluded to be reworked. Based on seismic and well-log data, clastic intercalations (probably of clays, silts and sands) within a halite-dominated sequence are thought to persist in the overlying Unit 6, although they diminish in thickness and frequency [START_REF] Gvirtzman | Intense salt deformation in the Levant Basin in the middle of the Messinian salinity crisis[END_REF][START_REF] Feng | Lithology and characteristics of the Messinian evaporite sequence of the deep Levant Basin, Eastern Mediterranean[END_REF]Meilijson et al., 2019). The expression of the end of the MSC is highly controversial. Until recently, the M-reflector of [START_REF] Ryan | Messinian badlands on the southeastern margin of the Mediterranean Sea[END_REF] (later renamed as the Top Erosion Surface, TES; Lofi et al., 2011aLofi et al., , 2011b) ) bounding Unit 6 at the top was [START_REF] Camerlenghi | Seismic markers of the Messinian salinity crisis in the deep Ionian Basin[END_REF]. Note how the MSC units are thinner, more difficult to distinguish and more deformed in the Sirt Abyssal Plain than in the IAP. PQ, UU and MU all onlap the structural highs of the Medina Ridge and VHS-2 sill. (b) High-resolution imaging of the lower part of the Plio-Quaternary (PQc unit) and upper part of the Messinian (UU) in the IAP [START_REF] Camerlenghi | Seismic markers of the Messinian salinity crisis in the deep Ionian Basin[END_REF]. The MSC-PQ boundary is a highly irregular surface, describing apparent V-shaped incisions (symbol V) of controversial origin (see subsection 4.6.1 for insights). Note the coherent deformation of PQc with the underlying MSC sequence and the absence of fluvial facies within the incisions (Unit II is made of lower Pliocene dolomitic marls recovered in Site 374 drilled nearby the seismic line; see text). (c) Multichannel seismic reflection profile MEM-07-203 running approximately parallel to the Malta Escarpment and showing the relationship between Unit 2 of [START_REF] Micallef | Evidence of the Zanclean megaflood in the eastern Mediterranean Basin[END_REF] with the overlying and underlying Zanclean and Messinian sediments, respectively (modified from [START_REF] Spatola | A single-stage megaflood at the termination of the Messinian salinity crisis: geophysical and modelling evidence from the eastern Mediterranean Basin[END_REF]. (d) Uninterpreted (left) and interpreted (right) seismic profiles showing the cyclic and channelized nature of the uppermost Messinian observed in the offshore Sirt Basin (modified from [START_REF] Bowman | A comprehensive review of the MSC facies and their origins in the offshore Sirt Basin, Libya[END_REF]. (e) Interpreted 2D regional WNW-ESE seismic profile crossing the continental shelf and offshore Levant Basin and the Herodotus Abyssal Plain [START_REF] Jagger | Tectono-stratigraphic evolution of the SE Mediterranean passive margin, offshore Egypt and Libya[END_REF]. Note the lateral continuity of the Messinian MU. (f) Seismic profile from the Levant Basin showing the 6 sub-units distinguished inside the MU as well as its lower (N-reflector) and upper (M-reflector) boundaries (modified from [START_REF] Gvirtzman | Intense salt deformation in the Levant Basin in the middle of the Messinian salinity crisis[END_REF]. (g) High-resolution seismic reflection image with wireline logs from Aphrodite-2 well illustrating that M-reflector previously considered as top evaporitic sequence and M/P boundary here consists of a ~100-m-thick unit (i.e. Unit 7 of Gvirtzman et al., 2017) in which different layers are distinguished (modified from [START_REF] Gvirtzman | Intra-Messinian truncation surface in the Levant Basin explained by subaqueous dissolution[END_REF]. (h) Interpreted and uninterpreted seismic profiles imaging the Mavqi'im and Afiq formations described in the canyons on the Levant continental margin (modified from Ben [START_REF] Ben Moshe | Estimating drawdown magnitudes of the Mediterranean Sea in the Levant basin during the Lago Mare stage of the Messinian Salinity Crisis[END_REF].

F. Andreetto et al. considered to be the Miocene/Pliocene boundary (Fig. 8f; [START_REF] Ryan | Messinian badlands on the southeastern margin of the Mediterranean Sea[END_REF][START_REF] Gvirtzman | Intense salt deformation in the Levant Basin in the middle of the Messinian salinity crisis[END_REF][START_REF] Feng | Lithology and characteristics of the Messinian evaporite sequence of the deep Levant Basin, Eastern Mediterranean[END_REF]. Instead, [START_REF] Gvirtzman | Intra-Messinian truncation surface in the Levant Basin explained by subaqueous dissolution[END_REF] showed that in higher resolution seismic data the M-reflector/ TES is a bundle of reflectors forming a distinct layer (Unit 7) overlying a truncation surface (i.e. Unit 6/7 boundary) that they re-labelled intra-Messinian truncation surface (IMTS; Fig. 8g) and ascribed to subaqueous dissolution rather than subaerial incision (e.g. [START_REF] Bertoni | Major erosion at the end of the Messinian Salinity Crisis: evidence from the Levant Basin, Eastern Mediterranean[END_REF]Lofi et al., 2011aLofi et al., , 2011b;;[START_REF] Kartveit | Evidence of sea level drawdown at the end of the Messinian salinity crisis and seismic investigation of the Nahr Menashe unit in the northern Levant Basin, offshore Lebanon[END_REF][START_REF] Madof | Discovery of vast fluvial deposits provides evidence for drawdown during the late Miocene Messinian salinity crisis[END_REF]. This conclusion was recently corroborated by the independent study of [START_REF] Kirkham | The demise of a 'salt giant' driven by uplift and thermal dissolution[END_REF]. Analysis of gamma-ray and resistivity logs in three deep basin wells (Aphrodite-2, Myra-1, Sara-1; Fig. 2b) and correlation with the Or-South-1 well (located between the deep basin and the shelf) showed that Unit 7 maintains a constant thickness of ~100 mthick and consists of clastic-rich anhydrite of undisclosed provenance. Meilijson et al. (2019)'s lithological interpretation of industrial boreholes slightly farther to the NE (Fig. 2b) give Unit 7 a significantly smaller thickness (5 m;Fig. 3b). Independent studies offshore Lebanon and Syria [START_REF] Kartveit | Evidence of sea level drawdown at the end of the Messinian salinity crisis and seismic investigation of the Nahr Menashe unit in the northern Levant Basin, offshore Lebanon[END_REF][START_REF] Madof | Discovery of vast fluvial deposits provides evidence for drawdown during the late Miocene Messinian salinity crisis[END_REF] describe a unit (Nahr Menashe complex) interpreted as a thicker (up to 300 m; [START_REF] Madof | Discovery of vast fluvial deposits provides evidence for drawdown during the late Miocene Messinian salinity crisis[END_REF], but lateral equivalent of [START_REF] Gvirtzman | Intra-Messinian truncation surface in the Levant Basin explained by subaqueous dissolution[END_REF]'s Unit 7. Based on its channelized morphology identified upslope near the Lebanese coast, [START_REF] Kartveit | Evidence of sea level drawdown at the end of the Messinian salinity crisis and seismic investigation of the Nahr Menashe unit in the northern Levant Basin, offshore Lebanon[END_REF] and [START_REF] Madof | Discovery of vast fluvial deposits provides evidence for drawdown during the late Miocene Messinian salinity crisis[END_REF] interpreted the Nahr Menashe Unit and the IMTS underneath as fluvial in origin, deposited/ formed on a subaerially exposed floor of the Levant Basin. Six [START_REF] Madof | Discovery of vast fluvial deposits provides evidence for drawdown during the late Miocene Messinian salinity crisis[END_REF] to eight [START_REF] Madof | Atlas of the Messinian Salinity Crisis markers in the Mediterranean and Black SeasMémoires de la Société géologique de France 179[END_REF] lobes were identified and are proposed to have stacked with precessional frequency. The Nahr Menashe sequence has been correlated by the same authors with the Abu Madi Fm. located within the Messinian canyons offshore Egypt (Abdel [START_REF] Aal | Tectonic evolution of the eastern Mediterranean Basin and its significance for hydrocarbon prospectivity in the ultradeepwater of the Nile Delta[END_REF][START_REF] Loncke | The Nile deep-sea fan: An example of interacting sedimentation, salt tectonics, and inherited subsalt paleotopographic features[END_REF][START_REF] Abdel-Fattah | Petrophysical characteristics of the messinian abu madi formation in the baltim east and north fields, offshore Nile delta, Egypt[END_REF], the Handere Formation offshore Turkey [START_REF] Radeff | Sedimentary evidence for late Messinian uplift of the SE margin of the Central Anatolian Plateau: Adana Basin, southern Turkey[END_REF] and with the Eosahabi deposits offshore Libya [START_REF] Bowman | A comprehensive review of the MSC facies and their origins in the offshore Sirt Basin, Libya[END_REF]. This interpretation implies a low base-level during the final stage of the MSC. [START_REF] Manzi | The onset of the Messinian salinity crisis in the deep Eastern Mediterranean basin[END_REF] and [START_REF] Meilijson | Deep-basin evidence resolves a 50-year-old debate and demonstrates synchronous onset of Messinian evaporite deposition in a non-desiccated Mediterranean[END_REF]Meilijson et al. ( , 2019) ) attempted astronomical dating of the abyssal MSC succession of the Levant Basin by integrating biostratigraphy on the pre-MU succession, reflector counting within the MU (only Meilijson et al., 2019) and well-log data (Fig. 3b). They achieved two contrasting results that gave rise to an outstanding controversy (Figs. 3b,4a). [START_REF] Manzi | The onset of the Messinian salinity crisis in the deep Eastern Mediterranean basin[END_REF] proposed that Stage 1 in the deep Levant is represented by a foraminifera-barren, evaporite-free shales interval labelled FBI (foraminifer barren interval) observed in the deep Aprodite-2 well and in the more proximal Myra-1 well. In this interpretation Unit 7 comprises the whole of Stage 3 (with the IMTS corresponding to the Stage 2/3 transition) and all halite deposition took place in ~50 kyr estimated during Stage 2 of the MSC (Fig. 1a; Roveri et al., 2014a). By contrast, the FBI is not present in the Dolphin well targeted by Meilijson et al. (2019), which is located in an intermediate position between the Aprodite-2 and Myra-1 wells studied by [START_REF] Manzi | The onset of the Messinian salinity crisis in the deep Eastern Mediterranean basin[END_REF]Fig. 2b). Instead, in the Dolphin well a relatively openmarine, foraminifera-rich sequence extends below the (conformable) base of the MU, placed in correspondence to a ~2 to 5.5 m-thick anhydrite/shale (Unit 0; [START_REF] Manzi | The onset of the Messinian salinity crisis in the deep Eastern Mediterranean basin[END_REF]Meilijson et al., 2018, respectively). Astronomical tuning of the ~33 cycles counted in the MU in the Dolphin well, which are assumed to be precessional, results in the Main Halite body (i.e. Unit 0-4 of [START_REF] Gvirtzman | Intense salt deformation in the Levant Basin in the middle of the Messinian salinity crisis[END_REF][START_REF] Manzi | The onset of the Messinian salinity crisis in the deep Eastern Mediterranean basin[END_REF] spanning MSC Stage 1 and 2, the Interbedded Evaporites/ Unit 5 covering substage 3.1 and the Argillaceous Evaporites/Unit 6-7 to encompass the Lago-Mare phase (Figs. 3b,4a). In this interpretation from Meilijson et al. (2019), halite deposition in the Levant Basin started in Stage 1 and persisted throughout the entire MSC, including Stage 3, during which more allochthonous material was delivered to the basin (Fig. 3b). [START_REF] Madof | Atlas of the Messinian Salinity Crisis markers in the Mediterranean and Black SeasMémoires de la Société géologique de France 179[END_REF] and [START_REF] Madof | Discovery of vast fluvial deposits provides evidence for drawdown during the late Miocene Messinian salinity crisis[END_REF] also attempted an astronomical tuning of the Nahr Menashe Unit, concluding that it spans throughout substage 3.2 and part of substage 3.1. Feng et al. (2016) claim, however, that the impressive thickness of clastics found in the Levantine MU is more indicative of distinct short-term events (shorter than the precession cycle) associated with transport of extraordinary power and magnitude.

Late Messinian sediments have also been recovered at several DSDP (129, 375, 376;Ryan et al., 1973;Hsü et al., 1978b) and ODP Sites (965, 967, 968;[START_REF] Emeis | Proc. ODP, Initial Reports[END_REF], but the assignment of the retrieved sediments to seismostratigraphic units is problematic. Nevertheless, they provide several key nuggets of precious information about the Stage paleoenvironment:

-Sites 965 and 966, located roughly on the crest of the Eratosthenes Seamount, just south of Cyprus (Fig. 2b), recovered soil structures above the evaporites indicating exposure above sea level (Robertson, 1998a(Robertson, , 1998b;;Maillard et al., 2011a;[START_REF] Reiche | The impact of salt on the late Messinian to recent tectonostratigraphic evolution of the Cyprus subduction zone[END_REF]. -ODP Sites 967 and 968, located at the base of the northern and southern slope of Eratosthenes Seamount (Fig. 2b), respectively, revealed the presence, within the MSC interval, of calcareous turbidites with Cyprus-derived clasts and clays containing Ammonia tepida, Cyprideis pannonica and pulmonate gastropods (Blanc-Valleron et al., 1998;Robertson, 1998a,b;[START_REF] Reiche | The impact of salt on the late Messinian to recent tectonostratigraphic evolution of the Cyprus subduction zone[END_REF]. -Abundant Cyprideis pannonica specimens were also recovered from DSDP Sites 375 and 376 drilled on the crest of the Florence Rise, west of Cyprus (Fig. 2b; [START_REF] Hsü | Initial reports of the deep sea drilling project: DSDP volume XLII Part[END_REF]. -Abundant, well-preserved Ammonia tests and Cyprideis specimens are also known from Site 129A (Fig. 2b), occurring with dwarf planktonic foraminifera (Ryan et al., 1973).

All the evidence listed above suggest that a base-level fall leading to subaerial exposure occurred at some point(s) during Stage 3 in the Eastern Mediterranean (Ryan, 2009). However, it must be kept in mind that both the Florence Rise and Eratosthenes Seamount are likely to have been much more elevated during the Messinian than they are today because of Pliocene-Quaternary subsidence related to the Cyprus subduction zone (Robertson, 1998a(Robertson, , 1998b;;Maillard et al., 2011a;[START_REF] Reiche | The impact of salt on the late Messinian to recent tectonostratigraphic evolution of the Cyprus subduction zone[END_REF].

Sites 375 and 376 display several discrete layers of primary and clastic gypsum interbedded in the Cyprideis-rich mudstones [START_REF] Mcculloch | Sr isotope constraints on the Mediterranean environment at the end of the Messinian salinity crisis[END_REF][START_REF] Lugli | The deep record of the Messinian salinity crisis: evidence of a non-desiccated Mediterranean Sea[END_REF]. This succession shares several similarities with sites drilled in the Western Mediterranean (e.g. ODP 654A) and Ionian Basin (DSDP 374), where they have been correlated with the seismic Upper Unit (see subsections 4. 5 and 4.6). This may suggest that a Western Mediterranean-like gypsum-bearing UU was also locally deposited in the easternmost abyss of the Mediterranean (see [START_REF] Günes | Internal seismic stratigraphy of the Messinian evaporites across the northern sector of the eastern Mediterranean Sea[END_REF].

Levant continental margin

Evaporitic and non-evaporitic deposits are buried beneath PQ deposits (Yafo Formation) along the Levant continental margin, where they are mostly preserved within canyons carved underwater in pre-Messinian time [START_REF] Druckman | The buried Afiq Canyon (eastern Mediterranean, Israel): a case study of a Tertiary submarine canyon exposed in Late Messinian times[END_REF][START_REF] Lugli | Evidence of clastic evaporites in the canyons of the Levant Basin (Israel): implications for the Messinian Salinity Crisis[END_REF]. Within the Afiq canyon, [START_REF] Druckman | The buried Afiq Canyon (eastern Mediterranean, Israel): a case study of a Tertiary submarine canyon exposed in Late Messinian times[END_REF] distinguished three formations in the Messinian sequence: the evaporitic Mavqi'im Formation, 115 mthick and mostly composed of anhydrite in places interbedded with micritic limestones; the Be'eri Formation, cmprising gypsum; the Afiq Formation, varying in thickness from 30 to 90 m and consisting of conglomerates, sandstones and marls interpreted as representing fluvial and lacustrine-marsh environments [START_REF] Druckman | The buried Afiq Canyon (eastern Mediterranean, Israel): a case study of a Tertiary submarine canyon exposed in Late Messinian times[END_REF]. The Afiq Fm. is only present in the deepest portions of the canyon where it overlies the Mavqi'im Fm. By contrast, the Be'eri gypsum is only found along the canyon shoulders covered by the Pliocene, at elevations > m with respect to the Mavqi'im Fm. Based on Sr values, [START_REF] Druckman | The buried Afiq Canyon (eastern Mediterranean, Israel): a case study of a Tertiary submarine canyon exposed in Late Messinian times[END_REF] attributed the Mavqi'im Fm. to MSC Stage 1, the Be'eri Fm. to substage 3.1 and the Afiq Fm. to the Lago-Mare phase. These authors also suggested that gypsum precipitation occurred under subaqueous conditions, with the water level ~600 m (i.e. the difference in altitude between the Mavqi'im and Be'eri fms.) higher during the deposition of the Be'eri Fm. Two base-level falls of approximately the same magnitude are thought to have occurred between the evaporitic phases and after

APPENDIX D

Mavqi'im deposition. A lowstand phase was therefore in force during Afiq deposition [START_REF] Druckman | The buried Afiq Canyon (eastern Mediterranean, Israel): a case study of a Tertiary submarine canyon exposed in Late Messinian times[END_REF].

However, combining stratigraphic, sedimentological and geochemical (Sr isotopes) considerations, [START_REF] Lugli | Evidence of clastic evaporites in the canyons of the Levant Basin (Israel): implications for the Messinian Salinity Crisis[END_REF] revealed the clastic nature of both the Mavqi'im and Be'eri evaporites and suggested the persistent drowning of the canyon(s), filled with turbidites [START_REF] Lugli | Evidence of clastic evaporites in the canyons of the Levant Basin (Israel): implications for the Messinian Salinity Crisis[END_REF]. Due to the presence of clastic evaporites, [START_REF] Gvirtzman | Intra-Messinian truncation surface in the Levant Basin explained by subaqueous dissolution[END_REF] suggested that the Mavqiim Formation is a condensed section encompassing MSC Stage 2 and early Stage 3, while the evaporite-free Afiq Formation represents the Lago-Mare phase.

Ben Moshe et al. ( 2020) ascribed (at least part of) the Afiq Fm. to the whole of Stage 3 as a wedge-shape clastic complex lying on top of the Mavqi'im Fm. and with the basal surface corresponding to the correlative conformity of the MES developed landward, at the expense of the Mavqi'im Fm (Fig. 8h). Ben Moshe et al. ( 2020) distinguished a lower sub-unit composed of clastic gypsum and with fore-stepping and downstepping internal geometry typical of progradational wedges, and an upper sub-unit containing anhydrite fragments and marls with Lago-Mare fauna (e.g. Cyprideis torosa; [START_REF] Rosenfeld | The sieve pores of Cyprideis torosa (Jones, 1850) from the Messinian Mavqi'im Formation in the coastal plain and continental shelf of Israel as an indicator of paleoenvironment[END_REF] and with seismic characteristics typical of a transgressive systems tract. Incisions are reported throughout the Afiq Fm. at different depths, while erosional surfaces bound both sub-units [START_REF] Ben Moshe | Estimating drawdown magnitudes of the Mediterranean Sea in the Levant basin during the Lago Mare stage of the Messinian Salinity Crisis[END_REF]. In particular, the surface capping the upper subunit and correlated to the M horizon or TES basinward (see subsection 4.7.1) shows dendritic drainage patterns of gullies and channels [START_REF] Ben Moshe | Estimating drawdown magnitudes of the Mediterranean Sea in the Levant basin during the Lago Mare stage of the Messinian Salinity Crisis[END_REF].

Ben Moshe et al. ( 2020) identify the variation of base level specifically during Stage 3 by analyzing the morphology of truncation surfaces bounding the Afiq Formation on the continental margin of the Levant Basin. This suggests high amplitude fluctuations of base-level in the order of one hundred meters, with development of subaerial erosion surfaces and the deposition of clastics and incision by fluvial drainage systems that occurred during the lowstand phases, while aggradational units (of unknown lithological nature) accumulated during the highstand phases. According to their analysis, base level dropped down to a maximum 535 m during Afiq deposition (i.e. below the maximum 430 m estimated paleodepth of the Sicily Sill; [START_REF] Garcia-Castellanos | Catastrophic flood of the Mediterranean after the Messinian salinity crisis[END_REF], implying hydrological disconnection between the Eastern and Western basins at various times during Stage 3. A regression to 615-885 m is proposed to have occurred at the top of the Afiq Fm., pre-dating the abrupt refilling at the base of the Zanclean (e.g. [START_REF] Micallef | Evidence of the Zanclean megaflood in the eastern Mediterranean Basin[END_REF][START_REF] Micallef | Geomorphic evolution of the Malta Escarpment and implications for the Messinian evaporative drawdown in the eastern Mediterranean Sea[END_REF][START_REF] Garcia-Castellanos | The Zanclean megaflood of the Mediterranean-Searching for independent evidence[END_REF][START_REF] Spatola | A single-stage megaflood at the termination of the Messinian salinity crisis: geophysical and modelling evidence from the eastern Mediterranean Basin[END_REF].

Summary of the offshore Stage 3 record

Knowledge of the Stage 3 sequence offshore is mainly based on the integration of seismic interpretations and analysis of material recovered from fragmentary and unevenly distributed DSDP/ODP/industrial cores.

▪ MSC sediments are absent on the eroded continental shelves bordering the deep basins, except in the Messinian thalwegs and at their mouth. Here the PQ lies directly above the MES which, in turn, cuts through the pre-MSC deposits (Fig. 7e). A similar stratigraphic arrangement is found along the middleupper slopes (Fig. 7d), although the presence of a thin, possibly incomplete UU in morphological depressions is sometimes postulated. Seamounts also lack MSC Stage 3 sediments and are strongly incised by the MES (Fig. 7g). ▪ The thick UU is widespread and roughly present everywhere in the abyssal plains from west of the Alborán volcanic arc to the eastern edge of the Ionian Basin (Fig. 2b). In the abyssal plains its seismic facies appears homogeneous, comprising parallel and relatively continuous high amplitude reflections (Figs. 7bc). The UU pinches out towards the foot of continental slopes and seamounts (Figs. 7b,, where it can be irregularly bedded or relatively well bedded (Lofi et al., 2011a(Lofi et al., , 2011b)).

The uppermost part of the Bedded Units (defined in depressions physically disconnected from the abyssal plains and, therefore, from the UU; e.g. CMD and Corsica Basin; [START_REF] Maillard | Record of the Messinian Salinity Crisis in the SW Mallorca area (Balearic Promontory, Spain)[END_REF][START_REF] Thinon | Seismic markers of the Messinian Salinity Crisis in an intermediate-depth basin: data for understanding the Neogene evolution of the Corsica Basin (Northern Tyrrhenian Sea)[END_REF]Raad et al., 2021) shows seismic features comparable to those of the UU. ▪ Drill Sites revealed that the reflections of relatively high amplitude in seismic profiles correspond to gypsum and mudstone alternations with sporadic intercalations of massive to cross-bedded sandstones. Some mudstone interbeds contain low-diversity assemblages of benthic organisms (ostracods, foraminifera and diatoms) indicative of shallow to neritic environments. Except for dwarfed forms of planktonic foraminifera and the monospecific nannofossil assemblages described by [START_REF] Castradori | Calcareous nannofossils in the basal Zanclean of the Eastern Mediterranean Sea: remarks on paleoceanography and sapropel formation[END_REF], the rest of planktonic foraminifera and nannofossils are largely regarded as reworked. ▪ The deep Levant Basin contains a ~1.8-2.0 km-thick MU (Figs. 8e-f), consisting of 6 to 7 seismic units depending on the resolution of the seismic employed. In high resolution seismic data, the lateral equivalent of part of the UU is identified as a ~100-m-thick, clastic-rich, anhydrite layer (Unit 7 of Gvirtzman 1207) offshore Israel, thickening to 300 m offshore Lebanon (Nahr Menashe complex, [START_REF] Madof | Discovery of vast fluvial deposits provides evidence for drawdown during the late Miocene Messinian salinity crisis[END_REF]. The Levant Basin still has major controversies concerning the timing of halite deposition (~50 kyr vs ~550 kyr; [START_REF] Manzi | The onset of the Messinian salinity crisis in the deep Eastern Mediterranean basin[END_REF]vs Meilijson et al., 2019), the origin of the clastic accumulations overlying the halite (fluvial vs subaqueous) and the presence or absence of gypsum-mudstone cycles. ▪ Apart from the halite flow-related deformation, the Miocene/ Pliocene boundary (i.e. UU/PQ transition) is conformable in intermediate (e.g. Balearic Promontory) and deep (WAB, EAB, Algero-Balearic, Liguro-Provençal, Tyrrhenian, Ionian and Levant) depocenters, while it shows signs of erosion on the shelf domain and along the upper-middle continental slopes and seamounts. Clear arguments of floor exposure at the M/P boundary are absent in all drill sites but 978A.

The paleontological perspective

Paleontological data have been used to define and identify Stage sediments, but have also been a source of profound contention over the interpretation of its paleoenvironmental and paleohydrological nature. Biotic groups impacted by the evolution of the gateways linking the Mediterranean with the Atlantic, Indian Ocean and Paratethys include marine species (e.g. foraminifera, calcareous nannofossils, fish) and brackish water-species (ostracods, fish, mollusks and dinocysts endemic or with affinity to species of the Paratethys region) that were unable to migrate when these corridors were closed, and terrestrial species (e.g. mammals) that, conversely, got across the gateway during periods of exposure (see [START_REF] Colombero | Late Messinian mollusks and vertebrates from Moncucco Torinese, north-western Italy. Paleoecological and paleoclimatological implications[END_REF]Booth-Rea et al., 2018;Mas et al., 2018b). Analysis of these faunal datasets provides key insights into likely gateway dimensions and the timing of their opening, restriction and closure (e.g. [START_REF] Palcu | Middle Miocene paleoenvironmental crises in Central Eurasia caused by changes in marine gateway configuration[END_REF]. Furthermore, they are a key constraint on the water sources likely to have been affecting the Mediterranean during MSC Stage 3.

Ostracods

Ostracods are by far the most prolific faunal group during Stage 3. Brackish species are known from both land sections and deep-sea cores across the whole Mediterranean (see Fig. 2b for sites and references; Fig. 9a). Two characteristic biofacies are commonly distinguished: Biofacies 1 [START_REF] Bonaduce | Paleoecological interpretation of the latest Messinian sediments from southern Sicily (Italy)[END_REF] or Cyprideis assemblage [START_REF] Iaccarino | Paleoenvironment of uppermost Messinian sequences in the western Mediterranean (Sites 974, 975, and 978[END_REF]) consists of an monospecific population of Cyprideis species or of an oligotypic population dominated by Cyprideis species alongside rare specimens of Tyrrhenocythere ruggierii, Loxoconcha kochi, Loxoconcha muelleri and Caspiocypris alta); Biofacies 2 (Bonaduce F. Andreetto et al. and Sgarrella, 1999) or Loxocorniculina djafarovi assemblage [START_REF] Iaccarino | Paleoenvironment of uppermost Messinian sequences in the western Mediterranean (Sites 974, 975, and 978[END_REF]) has a higher species diversity characterized by the abundant occurrence of truly Paratethyan species belonging to the genera Amnicythere, Loxoconcha, Loxocauda, Cytheromorpha, Cyprinotus and Tyrrenhocythere (see species names in Fig. 9a). The number of species reported in the onshore sections is variable, ranging from a dozen (e. g. [START_REF] Caruso | The late Messinian "Lago-Mare" event and the Zanclean Reflooding in the Mediterranean Sea: new insights from the Cuevas del Almanzora section (Vera Basin, South-Eastern Spain)[END_REF] to more than sixty (e.g. [START_REF] Gliozzi | Paratethyan ostracod immigrants in Italy during the Late Miocene[END_REF][START_REF] Grossi | Late Messinian Lago-Mare ostracods and paleoenvironments of the central and eastern Mediterranean Basin[END_REF]. This variability is not explained, but it may result from the application of different taxonomic concepts that resulted in the recognition of more or fewer species (Stoica et al., 2016) or from local environmental conditions that differed from basin to basin and resulted in different patterns of colonization.

Compared to the onshore domain, the ostracod fauna offshore is impoverished. Monospecific assemblages of Cyprideis sp. (Sites 372, 129A, 376, 654A, 967, 968;Ryan et al., 1973;[START_REF] Cita | Messinian Paleoenvironments: New Evidence from the Tyrrhenian Sea (ODP LEG 107)[END_REF] or oligospecific assemblages dominated by Cyprideis and rare specimens of et al., 2016, Cosentino et al., 2018[START_REF] Sciuto | New reports of Paratethyan ostracods affinity from the Mediterranean Basin (Sicily, Italy)[END_REF]. (b) Photomicrographs of the Paratethyan dinoflagellate cyst Galeacysta etrusca under the optical microscope (left) and SEM (right) (modified from Do [START_REF] Couto | Lago Mare and the Messinian salinity crisis: evidences from the Alboran Sea (S. Spain)[END_REF][START_REF] Andreetto | Migration of the dinoflagellate Galeacysta etrusca and its implications for the Messinian Salinity Crisis[END_REF]. Scale=20 μm. (c) SEM microphotographs of the euryhaline, shallow-water benthic foraminifera Ammonia beccarii (1-spiral side, 2umbilical side) and Ammonia tepida (3-spiral side;Carnevale et al., 2019) and of the dwarf fauna of planktonic foraminifera from the Bajo Segura Basin (4; Corbí and Soria, 2016). (d) Photographs in polarized light (crossed nicols) of some specimens of Ceratolithus acutus (1)(2)[START_REF] Lein | Isotopic composition of sulfur compounds in sediments of the Pacific[END_REF] described in the Lago-Mare unit of Malaga (1-Do Couto et al., 2014), the Zorreras Mb. of Sorbas (2-Clauzon et al., 2015) and the Colombacci Fm. of the Northern Apennines (3-Popescu et al., 2017) and of destroyed (4) and intact ( 5) ascidian spicules of Micrascidiscus sp. [START_REF] Golovina | A Late Maeotian age (6.7-6.3 Ma) for the enigmatic "Pebbly Breccia" unit in DSDP Hole 380A of the Black Sea[END_REF]. Note that C. acutus specimens closely resemble ascidian spicules of Micrascidiscus sp., which may lead to misinterpretation of the C. acutus (see [START_REF] Golovina | A Late Maeotian age (6.7-6.3 Ma) for the enigmatic "Pebbly Breccia" unit in DSDP Hole 380A of the Black Sea[END_REF], and that pictures and 2 are identical, despite they are attributed to samples taken from two different localities. (e) Articulated skeletons of marine fish from substage 3.1 mudstone horizons in Cyprus (1-Aphanius crassicaudus; Manzi et al., 2016a) and substage 3.2 deposits in Cava Serredi (2-Mugil cf. cephalus; [START_REF] Carnevale | Fossil marine fishes and the "Lago Mare" event: Has the Mediterranean ever transformed into a brackish lake?[END_REF].

F. Andreetto et al. Candona sp. (Hole 974B, Iaccarino and Bossio, 1999) and L. muelleri (Hole 978, [START_REF] Iaccarino | Paleoenvironment of uppermost Messinian sequences in the western Mediterranean (Sites 974, 975, and 978[END_REF] are the more widely reported. Interestingly, these assemblages are always associated with Ammonia sp. tests and, in some cases, with other species of shallow-water, euryhaline benthic foraminifera (see subsection 5.4). Only in Hole 975, close to the M/P boundary is a more heterotypic ostracod assemblage found (Euxinocythere praebaquana, Amnicythere idonea, Leptocythere limbate, Loxocorniculina djafarovi, Candona sp., and Cyprideis sp.; [START_REF] Iaccarino | Paleoenvironment of uppermost Messinian sequences in the western Mediterranean (Sites 974, 975, and 978[END_REF] and lacking of euryhaline benthic foraminifera. The likely cause of the widespread barrenness of ostracods in most of the offshore samples is perhaps because environmental conditions in the deep basins (depth and/or salinity) where not suitable to permit population by this benthic fauna (see below for the ecological requirements; e.g. [START_REF] Hsü | Initial reports of the deep sea drilling project: DSDP volume XLII Part[END_REF] in reference to Site 374). Finally, one must bear in mind that studying these organisms require much more material (some hundreds of grams) than the quantity of core sediments usually processed (i.e. ~10 cm 3 ; Iaccarino and Bossio, 1999).

The paleoecology (salinity and depth ranges) of Stage 3 ostracods has been based on both observations of few species that still live in the Caspian and Black seas today and have affinities with the Stage 3 species and on the interpretation of sedimentological, geochemical and mineralogical data of the surrounding sediments (see [START_REF] Gliozzi | A late Messinian brackish water ostracod fauna of Paratethyan aspect from Le Vicenne Basin (Abruzzi, central Apennines, Italy)[END_REF]Grossi, 2008 andGrossi et al., 2008 for insights). Biofacies 1 is thought to represent very shallow water environments (i.e. <15 m) with salinity fluctuating between mesohaline and hypersaline when the euryhaline Cyprideis is dominant. Instead, more stable oligo-mesohaline water is inferred when the other species are more abundant in Biofacies 1. The variegated Biofacies 2, on the other hand, is thought to represent somewhat deeper environments (up to 100 m) and less salty conditions (oligo-low mesohaline; [START_REF] Gliozzi | Late Messinian lago-mare ostracod paleoecology: a correspondence analysis aroach[END_REF][START_REF] Grossi | Late Messinian Lago-Mare ostracods and paleoenvironments of the central and eastern Mediterranean Basin[END_REF][START_REF] Caruso | The late Messinian "Lago-Mare" event and the Zanclean Reflooding in the Mediterranean Sea: new insights from the Cuevas del Almanzora section (Vera Basin, South-Eastern Spain)[END_REF].

Some authors consider the time when the Paratethyan ostracods arrived in the Mediterranean to be well constrained (e.g. Roveri et al., 2008a;Grossi et al., 2011;Cosentino et al., 2018) by the scarce occurrence of the first Paratethyan immigrant Loxoconcha muelleri 20 cm below the ash layer in the Colla di Votta section, which has a 238 U-206 Pb age of 5.5320±0.0074 Ma [START_REF] Cosentino | Refining the Mediterranean "Messinian gap" with high-precision U-Pb zircon geochronology, central and northern Italy[END_REF], and in the chaotic deposits of the Adana Basin, ascribed to Stage 2 [START_REF] Faranda | Messinian paleoenvironmental changes in the easternmost Mediterranean Basin: Adana Basin, southern Turkey[END_REF]. Instead, the first appearance of Loxocorniculina djafarovi has been considered to coincide with the biofacies 1-2 shift and to have occurred Mediterranean-wide synchronously at 5.40 Ma (Roveri et al., 2008a;Grossi et al., 2011;[START_REF] Cosentino | Refining the Mediterranean "Messinian gap" with high-precision U-Pb zircon geochronology, central and northern Italy[END_REF]. Roveri et al. (2008a) also showed Biofacies 2 diversity as increasing linearly through the Lago-Mare phase, reaching its maximum diversity just beneath the Miocene/Pliocene boundary and before disappearing in the Pliocene. Following the claimed synchronicity of the FO of both Loxoconcha muelleri and Loxocorniculina djafarovi, Roveri et al. (2008a) and Grossi et al. (2011) recognized one biozone in each biofacies: the Loxoconcha muelleri Biozone, spanning from 5.59 to 5.40 Ma, and the Loxocorniculina djafarovi Biozone, whose boundaries correspond respectively to the first (5.40 Ma) and last occurrence (5.33 Ma) of L. djafarovi in the Mediterranean. This biozonation, erected by Grossi et al. (2011), is often used for dating incomplete successions (e.g. Vera Basin; Stoica et al., 2016;[START_REF] Caruso | The late Messinian "Lago-Mare" event and the Zanclean Reflooding in the Mediterranean Sea: new insights from the Cuevas del Almanzora section (Vera Basin, South-Eastern Spain)[END_REF]. However, the first appearance of a diversified ostracod assemblage (including Loxocorniculina djafarovi) occurred in already cycle 3 of the Sicilian Upper Gypsum at Eraclea Minoa (Fig. 3a; Grossi et al., 2015), which has an astronomical age of 5.45 Ma (Van der Laan et al., 2006) or 5.47 Ma [START_REF] Manzi | A new facies model for the Upper Gypsum of Sicily (Italy): chronological and paleoenvironmental constraits for the Messinian salinity crisis in the Mediterranean[END_REF]. Furthermore, the sudden appearance of Biofacies 2 and its linear, upward increase in diversity have not been recognized in localities like Nijar and Malaga, where biofacies 1 and 2 are found stacked in more than one lithological (possibly precession-controlled) cycle in the Lago-Mare succession [START_REF] Bassetti | Evolution of depositional environments after the end of Messinian Salinity Crisis in Nijar basin (SE Betic Cordillera)[END_REF][START_REF] Guerra-Merchán | Messinian Lago-Mare deposits near the strait of Gibraltar (Malaga basin, S Spain)[END_REF]. These findings argue that the appearance of Paratethyan ostracods in the Mediterranean may not have been synchronous, therefore casting serious doubts upon the biostratigraphic relevance of the Mediterranean ostracods.

Except for Cyprideis specimens, where species attribution is debated (see discussion in Stoica et al., 2016), the affinity of all other ostracod species observed in Mediterranean Stage 3 sediments (Fig. 9a) with those of the Eastern Paratethys basins (i.e. Dacian, Euxinic and Caspian) has been demonstrated in several publications (e.g. Ruggieri, 1967;[START_REF] Gliozzi | Paratethyan ostracod immigrants in Italy during the Late Miocene[END_REF]Stoica et al., 2016;[START_REF] Sciuto | New reports of Paratethyan ostracods affinity from the Mediterranean Basin (Sicily, Italy)[END_REF]. Only [START_REF] Bassetti | Ostracod faunas and brackish-water environments of the late Messinian Sapigno section (northern Apennines, Italy)[END_REF][START_REF] Bassetti | Evolution of depositional environments after the end of Messinian Salinity Crisis in Nijar basin (SE Betic Cordillera)[END_REF] have questioned the Paratethyan affinity by suggesting that species from the Northern Apennines and Nijar Basin have ambiguous affinities with Paratethyan fauna as described in the mainly Russian literature from the '60-'70s. However, these differences between the late Messinian Mediterranean and Paratethyan ostracods resulted from misidentifications and/or a different use of species nomenclature (Stoica et al., 2016). Recently acquired knowledge of the Pontian assemblages of the Dacian, Euxinic and Caspian basins now permit to trace the provenance of Mediterranean Stage 3 ostracod species from the entire Black Sea region (Stoica et al., 2016) and, for a few species, from the Dacian [START_REF] Stoica | Paleoenvironmental evolution of the East Carpathian foredeep during the late Miocene-early Pliocene (Dacian Basin; Romania)[END_REF][START_REF] Lazarev | From Khersonian drying to Pontian "flooding": Late Miocene stratigraphy and paleoenvironmental evolution of the Dacian Basin (Eastern Paratethys)[END_REF], Caspian [START_REF] Van Baak | Mediterranean-Paratethys connectivity during the Messinian salinity crisis: The Pontian of Azerbaijan[END_REF] and North Aegean (see references in Krijgsman et al., 2020a) basins.

The means by which the ostracods travelled from the Paratethys to and across the Mediterranean during Stage 3 is as crucial for reconstructing the Stage 3 paleoenvironment as it is poorly addressed in onshore studies or overlooked in seismic and computational studies. Two migratory mechanisms have been suggested:

1) the aerial dispersion of ostracods through the migration of aquatic birds [START_REF] Benson | The paleoecology of the ostracods of DSDP Leg 42A[END_REF][START_REF] Caruso | The late Messinian "Lago-Mare" event and the Zanclean Reflooding in the Mediterranean Sea: new insights from the Cuevas del Almanzora section (Vera Basin, South-Eastern Spain)[END_REF]; this hypothesis was proposed because, in a Mediterranean concluded to have been desiccated, it was the only possible migration mechanism.

2) direct aqueous migration by the ostracods themselves (which are planktonic in the larval stage) through the establishment of similar paleoenvironmental conditions; by this mechanism, the dispersion of Paratethyan ostracod fauna from right across the Mediterranean requires E-W intraconnection and a Mediterranean water-level high enough to reach the marginal basins [START_REF] Gliozzi | Paratethyan ostracod immigrants in Italy during the Late Miocene[END_REF]Stoica et al., 2016;[START_REF] Sciuto | New reports of Paratethyan ostracods affinity from the Mediterranean Basin (Sicily, Italy)[END_REF][START_REF] Sciuto | Full restoration of marine conditions after the late Messinian Mediterranean Lago-Mare phase in Licodia Eubea and Villafranca Tirrena areas (east Sicily)[END_REF].

Finally, Carnevale et al. (2006aCarnevale et al. ( , 2006b[START_REF] Carnevale | Did the Mediterranean marine reflooding precede the Miocene/Pliocene boundary? Paleontological and geochemical evidence from upper Messinian sequences of Tuscany, Italy. Paleogeogr[END_REF][START_REF] Carnevale | Fossil marine fishes and the "Lago Mare" event: Has the Mediterranean ever transformed into a brackish lake?[END_REF] recognized the Paratethyan affinity of the Mediterranean Stage 3 species but, in view of their occurrence with in-situ species of marine fish, they suggested that Stage 3 ostracods descended from a Paratethyan stock that migrated into the Mediterranean well before the MSC and survived the extreme salinity conditions of Stage 1 and 2 in marginal, fresher water refugia. In this scenario the brackish water ostracod assemblages found in Stage have no paleoecological significance for Stage 3 paleoenvironment (Carnevale et al., 2006a(Carnevale et al., , 2006b[START_REF] Carnevale | Did the Mediterranean marine reflooding precede the Miocene/Pliocene boundary? Paleontological and geochemical evidence from upper Messinian sequences of Tuscany, Italy. Paleogeogr[END_REF][START_REF] Carnevale | Fossil marine fishes and the "Lago Mare" event: Has the Mediterranean ever transformed into a brackish lake?[END_REF]. However, there are two, unflagged problems with this hypothesis: 1) the Mediterranean-Central Paratethys connection through the Trans-Tethyan gateway in Slovenia already closed in the early Tortonian [START_REF] Kováč | Badenian evolution of the Central Paratethys Sea: paleogeography, climate and eustatic sea-level changes[END_REF][START_REF] Sant | Changing seas in the Early-Middle Miocene of Central Europe: a Mediterranean approach to Paratethyan stratigraphy[END_REF][START_REF] Palcu | Middle Miocene paleoenvironmental crises in Central Eurasia caused by changes in marine gateway configuration[END_REF]; 2) No Paratethyan ostracod species have been found in the Mediterranean before the MSC (see [START_REF] Gliozzi | Paratethyan ostracod immigrants in Italy during the Late Miocene[END_REF].

Dinoflagellate cysts

Dinoflagellate cysts (dinocysts) are the fossil remains of unicellular protists that live in the upper water column of many water bodies (e.g. [START_REF] Zonneveld | Atlas of modern dinoflagellate cyst distribution based on 2405 data points[END_REF][START_REF] Mudie | Atlas of modern dinoflagellate cyst distributions in the Black Sea Corridor: from Aegean to Aral Seas, including Marmara, Black, Azov and Caspian Seas[END_REF]. They can be used as paleoenvironmental indicators and for biostratigraphy, providing the ages of speciation and extinction events, as well as supplying evidence of age diagnostic dispersals of characteristic taxa/assemblages. Influxes of these microorganisms into a basin may occur as the result of interconnection with another basin and dinocysts can therefore be useful indicators of the open gateways between adjacent basins and the resultant changes in conditions (e.g. [START_REF] Andreetto | Migration of the dinoflagellate Galeacysta etrusca and its implications for the Messinian Salinity Crisis[END_REF]. In the case of the MSC, presence of in situ marine and/or Paratethys dinocyst assemblages in a marginal basin are likely to indicate the presence of Atlantic and/or Eastern Paratethys water (respectively) in the Mediterranean and (relatively) high water level conditions (e.g. Pellen et al., 2017).

Palynological studies on the late Messinian Mediterranean dinocysts

APPENDIX D

record are rather scarce, confined to a limited number of outcrops (Malaga Basin, Do Couto et al., 2014;Northern Apennines, Bertini, 2006;Popescu et al., 2007;[START_REF] Iaccarino | The Trave section (Monte dei Corvi, Ancona, central Italy): an integrated paleontological study of the Messinian deposits[END_REF][START_REF] Cosentino | Orbitally forced paleoenvironmental and paleoclimate changes in the late postevaporitic Messinian of the central Mediterranean Basin[END_REF]Pellen et al., 2017;Caltanissetta Basin, Londeix et al., 2007) and deep wells (976B, 977A, 978A and 134B, [START_REF] Popescu | Lago Mare episodes around the Messinian-Zanclean boundary in the deep southwestern Mediterranean[END_REF]. These studies describe substage 3.1 as being barren of dinocysts. By contrast, substage 3.2 dinocyst assemblages are diverse particularly a few meters/tens of meters below the Miocene/Pliocene boundary and show recurrent vertical variation in abundance between brackish, Paratethyan-type taxa and marine stenohaline and euryhaline species. Taxa with Paratethyan affinities are largely considered to be autochthonous by all aforementioned authors. The extent to which reworking may have affected the marine assemblages is more controversial and debated between none (in Malaga and in the Apennines; Popescu et al., 2007;[START_REF] Couto | Lago Mare and the Messinian salinity crisis: evidences from the Alboran Sea (S. Spain)[END_REF]Pellen et al., 2017), partial (in the uppermost part of the Sicilian Upper Gypsum; [START_REF] Londeix | Messinian paleoenvironments and hydrology in Sicily (Italy): the dinoflagellate cyst record[END_REF] and total (in the Apennines; e.g. [START_REF] Bertini | The Northern Apennines palynological record as a contribute for the reconstruction of the Messinian paleoenvironments[END_REF][START_REF] Iaccarino | The Trave section (Monte dei Corvi, Ancona, central Italy): an integrated paleontological study of the Messinian deposits[END_REF][START_REF] Cosentino | Orbitally forced paleoenvironmental and paleoclimate changes in the late postevaporitic Messinian of the central Mediterranean Basin[END_REF]. Given the extent of the implications (i.e. re-establishment of a Mediterranean-Atlantic flow or connection earlier than the Zanclean; e.g. Pellen et al., 2017), this is an issue that will require further clarification.

A key dinocyst influencing our understanding of the late Miocene Lago-Mare phase is Galeacysta etrusca (Fig. 9b; see [START_REF] Andreetto | Biostratigraphic and paleoecological significance of Galeacysta etrusca in the "Lago-Mare" facies from the Mediterranean area (Neogene)[END_REF][START_REF] Popescu | Galeacysta etrusca complex: dinoflagellate cyst marker of Paratethyan influxes to the Mediterranean Sea before and after the peak of the Messinian Salinity Crisis[END_REF][START_REF] Andreetto | Migration of the dinoflagellate Galeacysta etrusca and its implications for the Messinian Salinity Crisis[END_REF] for more insights). This species was originally described from sediments in the Mediterranean [START_REF] Corradini | Etude des dinokystes a la limite Messinien-Pliocene dans la coupe Cava Serredi, Tos cane, Italie[END_REF]), but has since been discovered in much older deposits in Paratethys (Magyar et al., 1999a(Magyar et al., , 1999b)). The earliest recorded occurrence of Galeacysta etrusca is in sediments from the Pannonian Basin dated at ~8 Ma (Magyar et al., 1999a(Magyar et al., , 1999b)). It subsequently dispersed throughout Paratethys at ~6 Ma and was present in the Black Sea throughout the MSC interval (Grothe et al., 2014[START_REF] Andreetto | Migration of the dinoflagellate Galeacysta etrusca and its implications for the Messinian Salinity Crisis[END_REF]. Despite a Mediterranean-Eastern Paratethys connection that is thought to have been established at ~6.1 Ma [START_REF] Krijgsman | Rise and fall of the Paratethys Sea during the Messinian Salinity Crisis[END_REF][START_REF] Van Baak | Mediterranean-Paratethys connectivity during the Messinian salinity crisis: The Pontian of Azerbaijan[END_REF]Grothe et al., 2020), G. etrusca is not found in the Mediterranean during MSC Stages 1, 2 and 3.1 (5.97-5.42 Ma;[START_REF] Bertini | The Northern Apennines palynological record as a contribute for the reconstruction of the Messinian paleoenvironments[END_REF][START_REF] Londeix | Messinian paleoenvironments and hydrology in Sicily (Italy): the dinoflagellate cyst record[END_REF], Manzi et al., 2007, Iaccarino et al., 2008, Gennari et al., 2013) and is only reported in the uppermost part of the Lago-Mare phase, very close to the transition to the Pliocene (e.g. [START_REF] Bertini | The Northern Apennines palynological record as a contribute for the reconstruction of the Messinian paleoenvironments[END_REF][START_REF] Londeix | Messinian paleoenvironments and hydrology in Sicily (Italy): the dinoflagellate cyst record[END_REF]Popescu et al., 2007;[START_REF] Iaccarino | The Trave section (Monte dei Corvi, Ancona, central Italy): an integrated paleontological study of the Messinian deposits[END_REF][START_REF] Cosentino | Orbitally forced paleoenvironmental and paleoclimate changes in the late postevaporitic Messinian of the central Mediterranean Basin[END_REF]Pellen et al., 2017). This implies that Galeacysta etrusca may have migrated from Paratethys into the Mediterranean after 5.42 Ma or that environmental conditions in the Mediterranean and in its marginal basins were only suitable for this species (and more generally the whole dinocysts Paratethyan contingent) to proliferate in the uppermost Messinian. Several authors report multiple occurrences of Galeacysta etrusca within the Zanclean (e.g. [START_REF] Clauzon | Influence of Mediterranean sea-level changes on the Dacic Basin (Eastern Paratethys) during the late Neogene: the Mediterranean Lago Mare facies deciphered[END_REF][START_REF] Londeix | Messinian paleoenvironments and hydrology in Sicily (Italy): the dinoflagellate cyst record[END_REF]Popescu et al., 2007[START_REF] Popescu | Lago Mare episodes around the Messinian-Zanclean boundary in the deep southwestern Mediterranean[END_REF][START_REF] Couto | Lago Mare and the Messinian salinity crisis: evidences from the Alboran Sea (S. Spain)[END_REF][START_REF] Clauzon | New insights on the Sorbas Basin (SE Spain): the onshore reference of the Messinian salinity crisis[END_REF], but these interpretations are based on the use of an alternative stratigraphic model for the MSC sections (Fig. 4b; see [START_REF] Andreetto | Migration of the dinoflagellate Galeacysta etrusca and its implications for the Messinian Salinity Crisis[END_REF] for details).

Diatoms

Among the fresh-brackish organisms found in Stage 3 sediments are also species of diatoms. To date (and to our knowledge), there are no onshore studies that have ever looked for these organisms. By contrast, two samples from DSDP Site 124 in the Algero-Balearic Basin (Fig. 2b) revealed the presence of littoral planktonic forms accompanied by brackish water, and even freshwater, euryhaline, benthonic, and epiphytic species in considerable numbers (Hajós, 1973). Diatoms of undisclosed paleoecological significance are also reported from the ~60 cm-thick mudstone bed between an anhydrite and halite bed found in the last core of Site 134 (Ryan et al., 1973). According to Hajós (1973) and Ryan (2009), the diatoms found in these drill cores attest to an extremely low salinity and a base level in the Balearic and Valencia basins below wave action. Further study of these indicative species and a wider distribution is required to apply this interpretation more generally.

Foraminifera

A reasonably diverse benthic and planktic foraminiferal assemblage containing no age-diagnostic taxa have been found co-occurring with the brackish Paratethyan fauna in both the onshore and offshore record throughout the Mediterranean (Fig. 2b for localities and references).

The benthic foraminifera assemblage is dominated by euryhaline representatives of the genus Ammonia, which today dwell in marginal marine (lagoons, estuaries, fjords and deltas) and lacustrine environments at depths < 50 m and tolerate salinities of up to 50‰ [START_REF] Milker | A taxonomic guide to modern benthic shelf foraminifera of the western Mediterranean Sea[END_REF][START_REF] Consorti | Insights on the paleoecology of Ammonia (Foraminifera, Rotalioidea) from Miocene carbonates of central and southern Apennines (Italy)[END_REF]. Ammonia tepida and Ammonia beccari (Fig. 9c) are by far the most abundant species in both onshore (see Fig. 2b for localities and references) and offshore (e.g. Site 968A, Blanc-Valleron et al., 1998;Sites 375, 376, 965-968, Orszag-Sperber, 2006) localities, where they co-occur with ostracods belonging to Biofacies 1. Other commonly occurring benthic euryhaline taxa are Elphidium sp., Cribroelphidium excavatum, Haynesina sp., Nonion sp., Quinqueloculina sp., Discorbis sp. and Trichohyalus sp., Brizalina dentellata, Bulimina echinate and Bolivina spp. (Ryan et al., 1973;Hsü et al., 1978a[START_REF] Hsü | Initial reports of the deep sea drilling project: DSDP volume XLII Part[END_REF][START_REF] Rouchy | Paleoenvironmental changes at the Messinian-Pliocene boundary in the eastern Mediterranean (southern Cyprus basins): significance of the Messinian Lago-Mare[END_REF][START_REF] Rouchy | Late Messinian to Early Pliocene paleoenvironmental changes in the Melilla Basin (NE Morocco) and their relation to Mediterranean evolution[END_REF][START_REF] Rouchy | The end of the Messinian salinity crisis: evidences from the Chelif Basin (Algeria)[END_REF][START_REF] Iaccarino | The Trave section (Monte dei Corvi, Ancona, central Italy): an integrated paleontological study of the Messinian deposits[END_REF][START_REF] Caruso | The late Messinian "Lago-Mare" event and the Zanclean Reflooding in the Mediterranean Sea: new insights from the Cuevas del Almanzora section (Vera Basin, South-Eastern Spain)[END_REF]. These species are frequently mixed with poorly preserved and older in age bathyal species (e.g. [START_REF] Caruso | The late Messinian "Lago-Mare" event and the Zanclean Reflooding in the Mediterranean Sea: new insights from the Cuevas del Almanzora section (Vera Basin, South-Eastern Spain)[END_REF].

Planktic foraminifera are represented both by species whose last occurrence pre-dates the MSC (e.g. Praeorbulina spp., Paragloborotalia partimlabiata, P. siakensis, Neogloquadrina atlantica praeatlantica, Globigerinoides subquadratus, Globorotalia saheliana, Globorotalia conomiozea, Acarinina sp., Hedbergella sp.) and by taxa with extended biostratigraphic ranges (e.g. Sphaeroidinellopsis seminulina, Turborotalita quinqueloba, Globorotalia miotumida, Globoturborotalita decoraperta, Neogloboquadrina acostaensis, Neogloboquadrina spp., Orbulina universa, Globigerinoides trilobus, Globigerinoides obliquus, Globorotalia scitula, Globigerina bulloides, G. Mediterranea and G. humerosa; see Fig. 2b for references).

The mixing of foraminifera species with different ecological and salinity requirements and the widespread agreement that the brackish Paratethyan fauna are autochthonous (see subsection 5.1) has always complicated the interpretation of the origin of the foraminiferal assemblages. Among the benthic species, Ammonia taxa and the other benthic euryhaline taxa are generally considered autochthonous because they are typically well-preserved and their ecological and salinity requirements could be compatible with those of the Paratethyan ostracods.

The habitat of these benthic foraminifera today in environments both influenced by and disconnected from the open ocean indicates that the Stage 3 sediments in which they occur were deposited in a shallowwater environment subject to salinity fluctuations (Caruso et al., and references therein), but they do not provide insights into the water provenance. By contrast, the poor preservation, older age and low diversity of the bathyal taxa strongly suggest that these species are reworked [START_REF] Bassetti | Evolution of depositional environments after the end of Messinian Salinity Crisis in Nijar basin (SE Betic Cordillera)[END_REF][START_REF] Iaccarino | The Trave section (Monte dei Corvi, Ancona, central Italy): an integrated paleontological study of the Messinian deposits[END_REF][START_REF] Caruso | The late Messinian "Lago-Mare" event and the Zanclean Reflooding in the Mediterranean Sea: new insights from the Cuevas del Almanzora section (Vera Basin, South-Eastern Spain)[END_REF]. Their mode of life is also incompatible with the shallower water elements of the faunal assemblage. The planktic species which went extinct before the MSC are also undoubtedly reworked [START_REF] Iaccarino | The Trave section (Monte dei Corvi, Ancona, central Italy): an integrated paleontological study of the Messinian deposits[END_REF][START_REF] Caruso | The late Messinian "Lago-Mare" event and the Zanclean Reflooding in the Mediterranean Sea: new insights from the Cuevas del Almanzora section (Vera Basin, South-Eastern Spain)[END_REF]. It is more challenging to discriminate between in situ and reworked specimens of the long range Neogene taxa. Most of them are considered to be reworked because of their scarcity, their occurrence with in-situ brackish organisms and their poor preservation (e.g. [START_REF] Iaccarino | The Trave section (Monte dei Corvi, Ancona, central Italy): an integrated paleontological study of the Messinian deposits[END_REF][START_REF] Caruso | The late Messinian "Lago-Mare" event and the Zanclean Reflooding in the Mediterranean Sea: new insights from the Cuevas del Almanzora section (Vera Basin, South-Eastern Spain)[END_REF]. A more complex controversy surrounds the long-range dwarf specimens (Fig. 9c) occurring in onshore substage 3.1 (di Tetto Fm. in the Trave section; [START_REF] Iaccarino | The Trave section (Monte dei Corvi, Ancona, central Italy): an integrated paleontological study of the Messinian deposits[END_REF] and Lago-Mare sediments (Upper Mb. of the Nijar Feos Fm., [START_REF] Fortuin | The Messinian of the Nijar basin (SE Spain): sedimentation, depositional environments and paleogeographic evolution[END_REF][START_REF] Aguirre | The Messinian post-evaporitic deposits of the Gafares area (Almeria-Nijar basin, SE Spain). A new view of the "Lago-Mare" facies[END_REF][START_REF] Bassetti | Evolution of depositional environments after the end of Messinian Salinity Crisis in Nijar basin (SE Betic Cordillera)[END_REF]Sorbas Basin, Roveri et al., 2019a;Bajo Segura Basin, Corbí and Soria, 2016;Colombacci Fm. in Northern Apennines localities, Casati et al., 1976;[START_REF] Colalongo | Paleoenvironmental study of the "Colombacci" Formation in Romagna (Italy): the cella section[END_REF][START_REF] Rio | Calcareous nannofossils (Monticino Quarry, Faenza)[END_REF] F. Andreetto et al. Popescu et al., 2007;Cyprus, Rouchy et al., 2001) and in some offshore localities (e.g. Sites 124, 125, 129A, 132, 134, 372, 376, 653, 974B, 975, 978;[START_REF] Cita | Inventory of biostratigraphical findings and problems[END_REF]Cita et al., 1978;[START_REF] Kastens | Proc. ODP, Init. Repts[END_REF][START_REF] Cita | Messinian Paleoenvironments: New Evidence from the Tyrrhenian Sea (ODP LEG 107)[END_REF][START_REF] Iaccarino | Paleoenvironment of uppermost Messinian sequences in the western Mediterranean (Sites 974, 975, and 978[END_REF]. This fauna is variably interpreted as:

1) reworked and size-sorted during transport, therefore lacking any paleoenvironmental significance (e.g. [START_REF] Kastens | Proc. ODP, Init. Repts[END_REF][START_REF] Iaccarino | Paleoenvironment of uppermost Messinian sequences in the western Mediterranean (Sites 974, 975, and 978[END_REF][START_REF] Fortuin | The Messinian of the Nijar basin (SE Spain): sedimentation, depositional environments and paleogeographic evolution[END_REF][START_REF] Bassetti | Evolution of depositional environments after the end of Messinian Salinity Crisis in Nijar basin (SE Betic Cordillera)[END_REF];

2) in situ and indicating normal marine conditions [START_REF] Aguirre | The Messinian post-evaporitic deposits of the Gafares area (Almeria-Nijar basin, SE Spain). A new view of the "Lago-Mare" facies[END_REF][START_REF] Braga | Testing models for the Messinian salinity crisis: the Messinian record in Almería, SE Spain[END_REF] or temporary Atlantic incursions [START_REF] Rouchy | Paleoenvironmental changes at the Messinian-Pliocene boundary in the eastern Mediterranean (southern Cyprus basins): significance of the Messinian Lago-Mare[END_REF];

3) in situ and indicative of high-stress environments [START_REF] Keller | Lilliput effect in late Maastrichtian planktic foraminifera: response to environmental stress: Paleogeography[END_REF], such as restricted and/or diluted marine environments (Corbí and Soria, 2016;Corbí et al., 2016[START_REF] Corbí | The step-by-step restriction of the Mediterranean (start, amplification, and consolidation phases) preceding the Messinian Salinity Crisis (climax phase) in the Bajo Segura basin[END_REF]. However, the paleoecological significance of dwarfism in foraminifer tests is not well understood and, given its potential implications for the Lago-Mare environment, it needs to be explored in greater detail.

Calcareous nannofossils and the C. acutus conundrum

Calcareous nannofossils are the fossil remains of coccolithophores, single-celled marine algae which dwell in the eutrophic and photic zone of the ocean (e.g. [START_REF] Ziveri | Present day coccolithophore-biogeography in the Atlantic Ocean[END_REF]. The potential recognition of marine calcareous nannofossils in marginal Stage 3 deposits would therefore have implications for the Mediterranean base-level and the hydrological riddle of MSC Stage 3. However, like foraminifera and dinocysts, the in situ versus reworking issue also impacts the nannoflora.

MSC Stage 3 is crossed by three important nannofossil bio-events astronomically calibrated in the ocean record: the top of Discoaster quinqueramus at 5.537 Ma, the base of Ceratolithus acutus at 5.36 Ma and the top of Triquetrorhabdulus rugosus at 5.231 Ma [START_REF] Backman | Biozonation and biochronology of Miocene through Pleistocene calcareous nannofossils from low and middle latitudes[END_REF][START_REF] Agnini | Calcareous nannofβossil biostratigraphy: historical background and alication in Cenozoic chronostratigraphy[END_REF]. Most of the (few) studies that addressed the nannoflora component of Stage 3 deposits did not report taxa belonging to the biozones defined by these bio-events, but only taxa of Cenozoic and Cretaceous age, clearly physically reworked (e.g. Sites 132, 134, 653, 654A, 967A, 969B, Ryan et al., 1973;[START_REF] Hsü | Initial reports of the deep sea drilling project: DSDP volume XLII Part[END_REF]Müller et al., 1990;[START_REF] Castradori | Calcareous nannofossils in the basal Zanclean of the Eastern Mediterranean Sea: remarks on paleoceanography and sapropel formation[END_REF]Piedmont Basin, Trenkwalder et al., 2008;[START_REF] Violanti | Micropaleontological analyses of the Narzole core: biostratigraphy and paleoenvironment of the late Messinian and early Zanclean of Piedmont (Northwestern Italy)[END_REF] Trave, Fonte dei Pulcini and Stingeti sections and Mondragone well in the Apennines, [START_REF] Cosentino | Environments and tectonic instability in central Italy (Garigliano Basin) during the late Messinian Lago-Mare episode: new data from the onshore Mondragone 1 well[END_REF][START_REF] Cosentino | Orbitally forced paleoenvironmental and paleoclimate changes in the late postevaporitic Messinian of the central Mediterranean Basin[END_REF]Cosentino et al., , 2018;;[START_REF] Iaccarino | The Trave section (Monte dei Corvi, Ancona, central Italy): an integrated paleontological study of the Messinian deposits[END_REF]). An exception is the nannoflora observed in the uppermost Messinian sediments at Sites 978A, 975B and 967A (Levant Basin; Fig. 2b). Here, among the plethora of reworked and long-ranging Neogene taxa, [START_REF] Castradori | Calcareous nannofossils in the basal Zanclean of the Eastern Mediterranean Sea: remarks on paleoceanography and sapropel formation[END_REF] reported the anomalous abundance of Sphenolithus spp (mostly Sphenolithus gr abies/moriformis). Although the assemblage points to the absence of a primary marine signature, the unlikely possibility that reworking and/or sorting lies behind the observed peak of Sphenolithus spp. led [START_REF] Castradori | Calcareous nannofossils in the basal Zanclean of the Eastern Mediterranean Sea: remarks on paleoceanography and sapropel formation[END_REF] to conclude that at least one incursion of marine water occurred during the (uppermost) Lago-Mare.

By contrast, some authors (i.e. Popescu et al., 2007[START_REF] Popescu | Lago Mare episodes around the Messinian-Zanclean boundary in the deep southwestern Mediterranean[END_REF][START_REF] Couto | Lago Mare and the Messinian salinity crisis: evidences from the Alboran Sea (S. Spain)[END_REF][START_REF] Clauzon | New insights on the Sorbas Basin (SE Spain): the onshore reference of the Messinian salinity crisis[END_REF]Pellen et al., 2017) described the nannofossil assemblage the Lago-Mare LM Unit in Malaga, the Zorreras Member in Sorbas, the uppermost di Tetto/Colombacci Fm. in some Apenninic localities and offshore in the Alborán Basin as having good preservation and showing no erratic fluctuations, all characteristics that led to their interpretation as autochthonous and to the conclusion that these sediments were deposited in a Mediterranean already replenished of Atlantic water (Fig. 4b). In addition, these authors reported the low abundance, but continuous presence of the biostratigraphic markers for the Zanclean Triquetrorhabdulus rugosus and Ceratolithus acutus (Fig. 9d) below the formally defined Miocene/Pliocene boundary [START_REF] Van Couvering | The base of the Zanclean Stage and of the Pliocene Series[END_REF] in several onshore and offshore Mediterranean (as wells as Paratethyan) localities (see [START_REF] Popescu | Ceratolithus acutus (= C. armatus), calcareous nannofossil marker of the marine reflooding that terminated the Messinian salinity crisis: Comment on "Paratethyan ostracods in the Spanish Lago-Mare: More evidence for interbasinal exchange at high Mediterranean sea level[END_REF] for details and a complete list of finding locations).

Such findings (especially that of C. acutus) are in sharp disagreement with most of the existing literature and have resulted in an important debate amongst the MSC community (e.g. Popescu et al., 2007, 2008vs Roveri et al., 2008cand Stoica et al., 2016vs Popescu et al., 2017), not only for their paleoenvironmental implications (i.e. presence of Atlantic water in the Mediterranean), but also for the chronostratigraphic repercussions (Fig. 4b). The chronostratigraphic value of C. acutus lies in its short temporal distribution straddling the M/P boundary (astrochronologically calibrated at 5.332 Ma;[START_REF] Van Couvering | The base of the Zanclean Stage and of the Pliocene Series[END_REF][START_REF] Lourens | The Neogene period[END_REF]. However, the corresponding biozone is established in oceanic areas (Zone CNPL1: 5.36-5.05 Ma;[START_REF] Backman | Biozonation and biochronology of Miocene through Pleistocene calcareous nannofossils from low and middle latitudes[END_REF][START_REF] Agnini | Calcareous nannofβossil biostratigraphy: historical background and alication in Cenozoic chronostratigraphy[END_REF] and is considered not applicable to the Mediterranean region during the MSC due to the harsh physicochemical conditions that are unsuitable for marine biota [START_REF] Di Stefano | Refinements of calcareous nannofossil biostratigraphy at the Miocene/Pliocene Boundary in the Mediterranean region[END_REF]. The interpretation of these nannofossil assemblages in the westernmost areas of the Mediterranean has been countered with several observations: (1) the observation of these age-diagnostic taxa is often not replicated by other studies (e.g. Roveri et al., 2008a;[START_REF] Van Baak | Messinian events in the Black Sea[END_REF]Krijgsman et al., 2020b) Krijgsman et al., 2020b). Recently, [START_REF] Golovina | A Late Maeotian age (6.7-6.3 Ma) for the enigmatic "Pebbly Breccia" unit in DSDP Hole 380A of the Black Sea[END_REF] showed that the morphology and size of C. acutus overlaps with the shape and dimensions of destroyed ascidian spicules (i.e. calcareous elements produced by benthic tunicates; Fig. 9d), providing an explanation for erroneous identification of C. acutus in the Black Sea Basin [START_REF] Golovina | A Late Maeotian age (6.7-6.3 Ma) for the enigmatic "Pebbly Breccia" unit in DSDP Hole 380A of the Black Sea[END_REF] and perhaps in the western Mediterranean Lago-Mare sediments as well.

Fish

Fossil fish remains provide information about salinity and depth and have been used to contradict the brackish nature of the Lago-Mare deposits by Carnevale et al. (2006aCarnevale et al. ( , 2006b[START_REF] Carnevale | Did the Mediterranean marine reflooding precede the Miocene/Pliocene boundary? Paleontological and geochemical evidence from upper Messinian sequences of Tuscany, Italy. Paleogeogr[END_REF][START_REF] Carnevale | Fossil marine fishes and the "Lago Mare" event: Has the Mediterranean ever transformed into a brackish lake?[END_REF] and [START_REF] Grunert | Estuarine Lago Mare fauna from the Tertiary Piedmont Basin indicates episodic Atlantic/Mediterranean exchange during the final stage of the Mediterranean Salinity Crisis[END_REF]. Euryhaline fish species inhabit marine to brackish environments and dominate settings with strong salinity variations while stenohaline fish have specific salinity requirements (marine, brackish, or freshwater) and cannot survive under different conditions. Demersal fish (i.e. those living in or immediately above the sea floor) have specific depth requirements, whereas pelagic fish occupy the water column within a wide range of depth variable from species to species. Fossil fish remains are found either as articulated or disarticulated skeletal parts, including teeth and otoliths, which are identified to the species level. Articulated fish skeletons typically indicate autochthonous deposition because of the difficulty in reworking and transporting intact skeletons. Otoliths and fish teeth are much more likely to be transported.

Otoliths and rare articulated skeletons (Fig. 9e) of marine and Paratethyan species have been reported from Stage 3 deposits, but commonly huge volumes of sediment are required to find even quite small numbers of these fossils (e.g. 20 tons from Moncucco, 6 tons from Cava Serredi, 700 kg from Capanne di Bronzo; [START_REF] Schwarzhans | Late Miocene-Early Pliocene evolution of Mediterranean gobies and their environmental and biogeographic significance[END_REF], much more than what is expected for normal marine deposits (i. e. < 30 kg; [START_REF] Agiadi | Connectivity controls on the late Miocene eastern Mediterranean fish fauna[END_REF]Karakitsios et al., 2017b). Substage 3.1 sediments contain articulated skeletons (Fig. 9e) of the marine fish species Lampanyctus licatae and Maurolicus muelleri, and the shallow water, euryhaline species Aphanius crassicaudus in the Lower Feos Member in the Nijar Basin (de la [START_REF] De La Chapelle | Decouverte de deux nouveaux gisements de poissons fossiles messiniens dans le bassin de Nijar-Carboneras (Andalousie Orientale): Signification paleoecologique et implications paleogeographiques[END_REF] and the marls of the first UG cycle in the Polemi Basin (Manzi et al., 2016a;Fig. 3a). Cava Serredi (Tuscany), Verduno and Moncucco (Piedmont) are the only other localities in which fish remains (only otoliths) in (claimed) substage 3.1 sediments are known (Carnevale et al., 2006a[START_REF] Carnevale | Did the Mediterranean marine reflooding precede the Miocene/Pliocene boundary? Paleontological and geochemical evidence from upper Messinian sequences of Tuscany, Italy. Paleogeogr[END_REF][START_REF] Carnevale | Fossil marine fishes and the "Lago Mare" event: Has the Mediterranean ever transformed into a brackish lake?[END_REF][START_REF] Grunert | Estuarine Lago Mare fauna from the Tertiary Piedmont Basin indicates episodic Atlantic/Mediterranean exchange during the final stage of the Mediterranean Salinity Crisis[END_REF].

The more diverse and abundant ichtyofaunal record occurs in substage 3.2 in a few marginal sections on the Italian peninsula (Ciabot Cagna in the Piedmont Basin; Cava Serredi and Podere Torricella in Tuscany; Capanne di Bronzo, La Vicenne and Ca' Ciuccio in thrust-top basins of the Northern and Central Apennines). The Lago-Mare fish F. Andreetto et al. remains mainly comprise otoliths of both euryhaline and stenohaline taxa indicative of marine, brackish, and freshwater habitats [START_REF] Carnevale | Fossil marine fishes and the "Lago Mare" event: Has the Mediterranean ever transformed into a brackish lake?[END_REF]. Three articulated skeletons of the euryhaline marine taxa Mugil cf. cephalus (Fig. 9e), the marine Indo-Pacific species Spratelloides gracilis and of Gobius sp. have been identified at Cava Serredi in a horizon < 1 m below the Miocene/Pliocene boundary (Carnevale et al., 2006b). The dominant stenohaline families in these assemblages are Gobiidae, a family of demersal fish occupying shallow-water marine, brackish and freshwater environments, and Myctophidae, which are marine mesopelagic fish that live below 200 m depth during the day, but feed at night in surface waters. A recent review of the Tortonian-Zanclean Gobiidae of the Mediterranean [START_REF] Schwarzhans | Late Miocene-Early Pliocene evolution of Mediterranean gobies and their environmental and biogeographic significance[END_REF] showed that the otoliths of this family, described by Carnevale et al. (2006a[START_REF] Carnevale | Did the Mediterranean marine reflooding precede the Miocene/Pliocene boundary? Paleontological and geochemical evidence from upper Messinian sequences of Tuscany, Italy. Paleogeogr[END_REF][START_REF] Carnevale | Fossil marine fishes and the "Lago Mare" event: Has the Mediterranean ever transformed into a brackish lake?[END_REF] and [START_REF] Grunert | Estuarine Lago Mare fauna from the Tertiary Piedmont Basin indicates episodic Atlantic/Mediterranean exchange during the final stage of the Mediterranean Salinity Crisis[END_REF] as belonging to marine Atlantic species, instead belong to brackish and freshwater species of Paratethyan affinity inhabiting sheltered prodelta environments. In fact, no normal marine demersal taxa were recognized in these assemblages by [START_REF] Schwarzhans | Late Miocene-Early Pliocene evolution of Mediterranean gobies and their environmental and biogeographic significance[END_REF]. As for the Myctophidae, the vast majority of the taxa belonging to this family were recovered in Moncucco and Verduno from alluvial plain silty mudstones along with terrestrial mammals (Dela Pierre et al., 2011;Colombero et al., 2017 and references therein), pointing to a physically reworked origin. When 87 Sr/ 86 Sr isotope ratios are measured [START_REF] Carnevale | Did the Mediterranean marine reflooding precede the Miocene/Pliocene boundary? Paleontological and geochemical evidence from upper Messinian sequences of Tuscany, Italy. Paleogeogr[END_REF][START_REF] Grunert | Estuarine Lago Mare fauna from the Tertiary Piedmont Basin indicates episodic Atlantic/Mediterranean exchange during the final stage of the Mediterranean Salinity Crisis[END_REF], the resulting Sr-based age of the otoliths is > 7 Ma, therefore further arguing against their in-situ origin. Since the good preservation of the otoliths suggests they did not suffer physical reworking (Carnevale et al., 2006a(Carnevale et al., , 2006b[START_REF] Carnevale | Did the Mediterranean marine reflooding precede the Miocene/Pliocene boundary? Paleontological and geochemical evidence from upper Messinian sequences of Tuscany, Italy. Paleogeogr[END_REF][START_REF] Carnevale | Fossil marine fishes and the "Lago Mare" event: Has the Mediterranean ever transformed into a brackish lake?[END_REF][START_REF] Grunert | Estuarine Lago Mare fauna from the Tertiary Piedmont Basin indicates episodic Atlantic/Mediterranean exchange during the final stage of the Mediterranean Salinity Crisis[END_REF], predators foraging in open marine settings and migrating to marginal environments are proposed as a way out of the enigma [START_REF] Carnevale | Did the Mediterranean marine reflooding precede the Miocene/Pliocene boundary? Paleontological and geochemical evidence from upper Messinian sequences of Tuscany, Italy. Paleogeogr[END_REF][START_REF] Carnevale | Fossil marine fishes and the "Lago Mare" event: Has the Mediterranean ever transformed into a brackish lake?[END_REF][START_REF] Grunert | Estuarine Lago Mare fauna from the Tertiary Piedmont Basin indicates episodic Atlantic/Mediterranean exchange during the final stage of the Mediterranean Salinity Crisis[END_REF][START_REF] Colombero | Late Messinian mollusks and vertebrates from Moncucco Torinese, north-western Italy. Paleoecological and paleoclimatological implications[END_REF]. However, Carnevale et al. (2006a) also rule out that so well preserved otoliths may have suffered post-mortem transport and action of the digestive acids in the stomach of predators. Rare findings of Myctophidae from Ciabot Cagna (3 species), Cava Serredi (1 species), Capanne di Bronzo (1 species) and Podere Torricella (6 species) [START_REF] Carnevale | Fossil marine fishes and the "Lago Mare" event: Has the Mediterranean ever transformed into a brackish lake?[END_REF] are all from sections where the host sediments have not been studied in sufficient detail to be clear about the in situ or reworked nature of the fossil assemblage. This lack of sedimentological uncertainty also extends to the stratigraphic position of many samples, because a stratigraphic log is provided for only a few sections (i.e. Ca' Ciuccio, Cava Serredi and Moncucco; Carnevale et al., 2006aCarnevale et al., , 2006b)). What this stratigraphic information suggests is that euryhaline fish taxa are widespread throughout substage 3.2, whereas strictly Myctophidae, which are an oceanic, marine stenohaline species, only occur very close to the base of the Pliocene, plausibly corresponding to the uppermost lithological cycle in substage 3.2 (~5.35-5.33 Ma;[START_REF] Carnevale | Fossil marine fishes and the "Lago Mare" event: Has the Mediterranean ever transformed into a brackish lake?[END_REF].

Summary of the Stage 3 paleontological record

The aquatic fossil record of MSC Stage 3 indicates that substage 3.1 in onshore sections is mostly barren, while diverse assemblages characterize substage 3.2 deposits. By contrast, the deep record as a whole contains relatively few, low diversity assemblages. This might be as a consequence either of the limited sample locations recovered from the offshore areas (see Fig. 2b) or because the environmental conditions in the intermediate-deep basins were less favorable for sustaining the life forms typical of the onshore domain. Nevertheless, the assemblages that are found in both marginal and deep locations comprise mixed brackish and marine species.

Brackish species are mostly represented by ostracods and dinocysts (and mollusks here not addressed because poorly studied; see [START_REF] Esu | Latest Messinian "Lago-Mare" Lymnocardiinae from Italy: close relations with the Pontian fauna from the Dacic Basin[END_REF][START_REF] Guerra-Merchán | Messinian Lago-Mare deposits near the strait of Gibraltar (Malaga basin, S Spain)[END_REF]. Prominent is the affinity of these late Messinian Mediterranean brackish species with the same species that were simultaneously dwelling in the Eastern Paratethyan basins (Dacian, Euxinic and Caspian) and in the North Aegean. Since these organisms were not present in the Mediterranean at any time before the MSC, they are considered, with a broad consensus, as in situ. This conclusion is further corroborated by the mixing of adult and juvenile forms in the ostracod assemblages and by the good preservation of the specimens, which do not show typical evidence of physical reworking like abrasion, dissolution, or fragmentation. Still problematic is the time of their arrival in the Mediterranean and their likelihood as biostratigraphic tool. From our review it seems more likely that truly Paratethyan species of ostracods entered the Mediterranean already during substage 3.1, when they colonized intermediate-deep settings, while they entered the marginal basins at different times during substage 3.2. As for dinocysts, characteristic is their occurrence only in the uppermost Messinian. However, it must be noted that samples from the substage 3.1 interval are rarely processed for dinocysts, especially in age modelequipped sections (Fig. 3a). The route followed by the Paratethyan immigrants is equally contested and important for paleoenvironmental and paleohydrological interpretations. In view of a desiccated Mediterranean, their migration can only have taken place passively by means of aquatic migratory birds. Conversely, the homogeneity of the ostracod assemblages throughout the Mediterranean marginal basins is more indicative of the presence of a water body fed by Eastern Paratethys and connecting all Mediterranean subbasins, therefore implying relatively high water-level conditions (at least at times when ostracod-bearing sediments deposited; see Andreetto et al., 2021).

Marine assemblages are composed by foraminifera, nannofossils, dinocysts and calcareous nannofossils. Their reworked or in situ nature is in many cases contested but critical for paleoenvironmental interpretation. The picture that emerges from our review is that an open marine signature is questionable in the foraminifera, nannofossils, dinocyst and fish records, as well as in other biotic groups (e.g. corals, echinoids and mammals) here not tackled (and for which we refer the reader to [START_REF] Dominici | Pliocene stratigraphic paleobiology in Tuscany and the fossil record of marine megafauna[END_REF]Carnevale et al., 2019). All marine representatives of the above mentioned categories were reintroduced into the Mediterranean only at the beginning of the Pliocene and at the expense of the Paratethys species that, instead, disappeared. Collectively, these observations lead us to conclude that the marine model as conceived by Carnevale et al. (2006aCarnevale et al. ( , 2006b[START_REF] Carnevale | Did the Mediterranean marine reflooding precede the Miocene/Pliocene boundary? Paleontological and geochemical evidence from upper Messinian sequences of Tuscany, Italy. Paleogeogr[END_REF][START_REF] Carnevale | Fossil marine fishes and the "Lago Mare" event: Has the Mediterranean ever transformed into a brackish lake?[END_REF] and [START_REF] Grunert | Estuarine Lago Mare fauna from the Tertiary Piedmont Basin indicates episodic Atlantic/Mediterranean exchange during the final stage of the Mediterranean Salinity Crisis[END_REF] has no foundation and therefore will not be further discussed.

The geochemical perspective

Variations in the water sources draining into the Mediterranean are expected to be reflected also in (geo)chemical properties of the paleodepositional environments. Important information about the nature of the connectivity framework of the Mediterranean can be gained by interpreting geochemical signals that respond to the presence or absence of an exchange with a chemically-unique water body. Four main geochemical proxies have been applied so far to MSC Stage 3 sedimentary and paleontological records. These includes both radiogenic (Sr isotope ratios) and stable isotopes (sulfate and oxygen) measured on fossils and minerals and hydrogen isotopes on molecular biomarkers. This section summarizes the dataset available for geochemical proxies (Fig. 10; Supplementary material 1) and its interpretation(s) for MSC Stage 3.

Strontium isotope ratios ( 87 Sr/ 86 Sr)

The available strontium isotope data for Stage 3 (Fig. 10a; Supplementary material 1) derive from measurements on both Ca-bearing fossils (ostracod valves, mollusk shells, fish otoliths; Fig. 9a) and minerals (calcite and gypsum), where Sr 2+ dissolved in an aqueous solution substitutes Ca atoms due to their similar ionic radius (e.g. [START_REF] Hajj | Radiogenic and "stable" strontium isotopes in provenance studies: a review and first results on archaeological wood from shipwrecks[END_REF]. Here we screen the available dataset and discuss only results that (1) reflect the original primary isotopic signal, i.e. the isotopic signal of the fluid at time of shell calcification or mineral precipitation, and (2) for which timing of mineral precipitation can be constrained. This screening excludes bulk carbonate samples (e.g. Colombacci limestones;

F. Andreetto et al. [START_REF] Bassetti | Paleoenvironmental significance of Messinian post-evaporitic lacustrine carbonates in the northern Apennines, Italy[END_REF], which contain carbonate compounds of various and/or unknown provenance, measurements from mollusk shells and otoliths (e.g. [START_REF] Carnevale | Did the Mediterranean marine reflooding precede the Miocene/Pliocene boundary? Paleontological and geochemical evidence from upper Messinian sequences of Tuscany, Italy. Paleogeogr[END_REF][START_REF] Grunert | Estuarine Lago Mare fauna from the Tertiary Piedmont Basin indicates episodic Atlantic/Mediterranean exchange during the final stage of the Mediterranean Salinity Crisis[END_REF]Roveri et al., 2019a), because they are made of mineral phases easily altered during diagenesis (e.g. aragonite; [START_REF] Marcano | Diagenetic incorporation of Sr into aragonitic bivalve shells: Implications for chronostratigraphic and paleoenvironmental interpretations[END_REF], and data coming from reworked material (e.g. all reworked gypsum or transported foraminifera). 87 Sr/ 86 Sr isotope ratios have also been measured by [START_REF] Müller | Origin and age of the Mediterranean Messinian evaporites: implications from Sr isotopes[END_REF] and Roveri et al. (2014b) on the halite beds recovered at Sites 134, 374 and 376 (Ryan et al., 1973;[START_REF] Hsü | Initial reports of the deep sea drilling project: DSDP volume XLII Part[END_REF]. Although they provide interesting interpretative aspects, we do not consider these Sr measurements because the position of Sr in the crystal lattice of halite is unknown and the removal of all contaminants, that is not a straightforward procedure (see Meilijson et al., 2019), is not clear it was achieved by [START_REF] Müller | Origin and age of the Mediterranean Messinian evaporites: implications from Sr isotopes[END_REF] and Roveri et al. (2014b). As a matter of fact, there is no consistency between data generated from roughly the same interval in Core 134 by Müller and Mueller (1991) (0.708968) and Roveri et al. (2014b) (0.708800-0.708896). Added to this is the uncertainty over the provenance of halite in Sites 134 and 374 (see subsections 4.2 and 4.6.1), which violates both criteria mentioned above.

The general trend of the Mediterranean 87 Sr/ 86 Sr isotope ratio during the MSC deviates from the ocean curve towards the less radiogenic values of the major peri-Mediterranean rivers and Paratethys and returns abruptly to oceanic values at the Miocene/Pliocene boundary (Fig. 10a inset). This trend is regarded to reflect the progressive restriction of Mediterranean-Atlantic exchange and the relative increase in the proportion of non-marine source waters (Topper et al., 2011;Roveri et al., 2014a). At first glance it seems that each MSC Stage was characterized by a well-defined range of Sr ratios (Fig. 10a inset), an observation that led Roveri et al. (2014b) to attribute a chronostratigraphic value to MSC 87 Sr/ 86 Sr ratios. A closer look, however, shows that MSC substages are anything but homogeneous with respect to 87 Sr/ 86 Sr ratios. At least in the marginal basins, local lithological differences in the catchments (each lithology carries a unique 87 Sr/ 86 Sr fingerprint; see subsection 8.1.1) explain the different Sr isotopic compositions from basin to basin (see [START_REF] Schildgen | Sea level and climate forcing of the Sr isotope composition of Late Miocene Mediterranean marine basins[END_REF][START_REF] Modestou | Precessional variability of 87 Sr/ 86 Sr in the late miocene sorbas basin: An interdisciplinary study of drivers of interbasin exchange[END_REF]Andreetto et al., 2021), therefore arguing against the use of 87 Sr/ 86 Sr ratios for chronostratigraphic purposes in the MSC record.

Most of the data characterizing substage 3.1 (Fig. 10a) are from the Eraclea Minoa gypsum (Fig. 5h). These data define a narrow range of Sr isotope ratios between 0.708747 and 0.708793 (García-Veigas et al., 2018). Similar values were reported from both Eraclea Minoa and the nearby Siculiana Marina section (0.708710-0.708760;Keogh and Butler, 1999;Fig. 5i). The dominance of Sicily samples gives the appearance of a consistent Sr isotope signal for gypsum beds. However, data points from elsewhere (Cyprus, Manzi et al., 2016a;DSDPs 122, 371 and 372 in the Algero-Balearic Basin, ODPs 652, 653 and 654 in the Tyrrhenian Basin, DSDP 374 in the Ionian Basin; Müller et al., 1990;[START_REF] Müller | Origin and age of the Mediterranean Messinian evaporites: implications from Sr isotopes[END_REF]Roveri et al., 2014b) display a wider range (from ~0.7087 to 0.708847; Fig. 10a) that may indicate a different hydrological regime for each basin (e.g. Müller et al., 1990;[START_REF] Müller | Origin and age of the Mediterranean Messinian evaporites: implications from Sr isotopes[END_REF]Ryan, 2009). The one published Sr isotope value for ostracods found within one of the marl interbeds at Eraclea Minoa also has a lower value outside the typical Sicily gypsum range (Grossi et al., 2015). This suggests that a different hydrological regime may also have characterised precession minima stages of the precessional cycle.

The Sr isotope dataset for the Lago-Mare phase includes the lowest values measured on MSC sediments (~0.7085 from between gypsum VI and VII at Eraclea Minoa; Fig. 3a; Grossi et al., 2015) and the widest range of ratios spanning from 0.7085 to 0.7091, which is above coeval oceanic values (Fig. 10a). Again, the conspicuously high Sr isotope values in substage 3.2 come from two areas, the marginal basins of southern Spain (Andreetto et al., 2021 and references therein;Figs. 5a -c) and the intermediate Polemi Basin on Cyprus [START_REF] Mcculloch | Sr isotope constraints on the Mediterranean environment at the end of the Messinian salinity crisis[END_REF]. The lower values are drawn from right across the intermediate-deep Mediterranean (Algero-Balearic, Sicily, Levant; Fig. 2a) and are therefore more likely to represent a Mediterranean-wide Sr isotope signal.

New Sr isotope data from Eastern Paratethys (i.e. Dacian and Caspian basins; Fig. 2b) are now available for the interval corresponding to MSC Stage 3 (inset Fig. 10a). The 87 Sr/ 86 Sr ratios of the Dacian Basin (0.708865-0.708982;[START_REF] Vasiliev | Strontium isotope ratios of the Eastern Paratethys during the Miocene/Pliocene transition; Implications for interbasinal connectivity[END_REF][START_REF] Grothe | The Messinian Salinity Crisis: a Paratethyan perspective[END_REF] are slightly lower than coeval ocean water (0.709020), but much higher than coeval Mediterranean values. However, the Dacian Basin is regarded as highly restricted from the Mediterranean throughout the MSC [START_REF] Vasiliev | Strontium isotope ratios of the Eastern Paratethys during the Miocene/Pliocene transition; Implications for interbasinal connectivity[END_REF]. By contrast, the Caspian has very low values (0.708402 to 0.708473, Grothe et al., 2020) which are thought to reflect both the very low Sr isotope ratio of the Volga river (0.708020;Vasiliev et al.,and references therein) and some input from the Mediterranean (Grothe et al., 2020).

Sulfate isotopes

Sulfur isotopic investigations have been carried out only on sulfate minerals (gypsum and more rarely anhydrite) of the MSC Stage 3 deposits with samples drawn from both onshore intermediate sequences (Caltanissetta Basin and Cypriot basins) and deep basinal records (Sites 122, 124, 125A, 132, 134, 372, 374, 375, 376, 652, 653, 654, 968, 969, 970;Fig. 10b;[START_REF] Fontes | Oxygen, Carbon, Sulfur and Hydrogen stable isotopes in carbonate and sulfate mineral phases of Neogene evaporites, sediments and in interstitial waters[END_REF][START_REF] Pierre | Contribution ~ l'Jtude sgdimentologique et isotopique des evaporites messiniennes de la Mediterrane; implications geodynamiques[END_REF]Pierre, , 1982;;[START_REF] Pierre | Isotope Composition of Messinian sediments from the Mediterranean Sea as indicators of paleoenvironments and diagenesis[END_REF][START_REF] Ricchiuto | Stable Isotopic investigation of Messinian sulfate samples from DSDP. LEG 42. Eastern Mediterranean Sea[END_REF]Pierre and Rouchy, 1990;Blanc-Valleron et al., 1998). Because the incorporation of dissolved sulfate into gypsum produces a nearly constant fractionation of δ 18 O (+3.5‰) and δ 34 S (+1.65‰) at earth surface temperatures (Thode and Monster, 1965;Lloyd, 1968;Warren, 2016), δ 18 O and δ 34 S isotopic values measured in gypsum should be corrected with the above mentioned fractionation factors in order to reconstruct the sulfate isotopic composition of the 87 Sr/ 86 Sr isotope data sourced from ostracod valves and gypsum crystals (see Supplementary material 1 and subsection 6.1 for references). Data are plotted with the global 87 Sr/ 86 Sr seawater curve (McArthur et al., 2012). Error bars indicate analytical error, which is so small in some cases that no error bars are visible at this scale. To not complicate the figure, horizontal error bars have not been added for the sections/cores unprovided of a chronostratigraphic framework and for which age uncertainties are present (i.e. all but Nijar and Vera basins, Eraclea Minoa and onshore Cyprus; see Fig. 3). Note that none of the 87 Sr/ 86 Sr isotope ratios but one from Nijar plot on the ocean curve. In the inset is shown the Mediterranean Sr record for the entire MSC as well as the time-equivalent Eastern Paratethys record (modified after Andreetto et al., 2021). (b) Plot of δ 34 S SO4 and δ 18 O SO4 in Stage 3 gypsum and anhydrite beds from onshore and offshore localities (see Supplementary material 1 and subsection 6.2 for references). No measures are available from the marginal basins, where gypsum did not deposit during Stage 3. The dark blue and black rectangles represent the sulfate isotopic composition of the Global Messinian ocean and Stage 1 (PLG) evaporites, respectively. The light blue area represents the sulfate isotopic composition of mixtures of Messinian marine waters with non-marine sources. The red area represents the isotopic composition of the residual sulfate ion in a basin where marine Messinian sulfate is consumed by microbial sulfate reduction to produce H 2 S. The arrow represents the isotope trajectory of dissolved sulfate resulting from the mixing of residual 34 Senriched sulfate produced by MSR and 34 S-depleted sulfate produced by H 2 S oxidation. All the published δ 34 S SO4 and δ 18 O SO4 values are provided corrected with the fractionation factors δ 34 S=+1.65‰ and δ 18 O=+3.5‰ to smooth the isotopic fractionation effects experienced by dissolved sulfate and to reason on values reproducing the isotopic composition at the time of gypsum precipitation. (c) δD isotopes of C 29 and C 31 n-alkanes and C 37 and C 38 long chain alkenones recorded in the Stage 3 gypsums and marls of the Eraclea Minoa section (modified from Vasiliev et al., 2017). Blue lines indicate the values recorded in the present day lacustrine settings for the n-alkanes [START_REF] Sachse | δD values of individual n-alkanes from terrestrial plants along a climatic gradient -implications for the sedimentary biomarker record[END_REF] and in the alkenones from the Mediterranean in the recent times [START_REF] Van Der Meer | Hydrogen isotopic compositions of long-chain alkenones record freshwater flooding of the Eastern Mediterranean at the onset of sapropel deposition[END_REF]. Error bars indicate standard errors of the mean.

F. Andreetto et al. basin waters at the time of gypsum formation.

The deep Mediterranean samples exhibit a wide range of δ 34 S SO4, but the majority of samples display δ 34 S SO4 values between 18 and 22‰, strongly indicative of a marine origin of the sulfate forming the gypsum (Fig. 10b; [START_REF] Fontes | Oxygen, Carbon, Sulfur and Hydrogen stable isotopes in carbonate and sulfate mineral phases of Neogene evaporites, sediments and in interstitial waters[END_REF][START_REF] Pierre | Contribution ~ l'Jtude sgdimentologique et isotopique des evaporites messiniennes de la Mediterrane; implications geodynamiques[END_REF]Pierre, , 1982;;[START_REF] Pierre | Isotope Composition of Messinian sediments from the Mediterranean Sea as indicators of paleoenvironments and diagenesis[END_REF]Pierre and Rouchy, 1990;Blanc-Valleron et al., 1998). The δ 34 S SO4 values lower than marine sulfate in the dataset are generally considered to represent a greater influence of continental sulfate input to the basin (Fig. 10b; [START_REF] Pierre | Contribution ~ l'Jtude sgdimentologique et isotopique des evaporites messiniennes de la Mediterrane; implications geodynamiques[END_REF][START_REF] Pierre | Isotope Composition of Messinian sediments from the Mediterranean Sea as indicators of paleoenvironments and diagenesis[END_REF]Pierre and Rouchy, 1990). By contrast, the data display δ 18 O SO4 isotopic values that deviate substantially from marine δ 18 O SO4 values towards higher values (Fig. 10). This is consistent with the influence of sulfate produced by reoxidation of reduced sulfur compounds generated by microbial sulfate reduction (MSR; [START_REF] Kaplan | Microbiological fractionation of sulphur isotopes[END_REF]Brunner and Bernasconi, 2005;Sim et al., 2011;Leavitt et al., 2013). The microbial use of SO 4

2-leads to an equilibration of δ 18 O SO4 with ambient water oxygen, whereas the δ 34 S SO4 returns towards its initial value as a higher fraction of sulfide produced by MSR is re-oxidated. This mechanism has been suggested for Sites in the Algero-Balearic, Tyrrhenian and Ionian basins and offshore Cyprus [START_REF] Pierre | Contribution ~ l'Jtude sgdimentologique et isotopique des evaporites messiniennes de la Mediterrane; implications geodynamiques[END_REF][START_REF] Pierre | Isotope Composition of Messinian sediments from the Mediterranean Sea as indicators of paleoenvironments and diagenesis[END_REF]Pierre and Rouchy, 1990). Although some authors have suggested that partial equilibration of sulfate oxygen toward δ 18 O H2O values of the basin enriched in heavy oxygen isotopes by evaporation have led to an increase in δ 18 O SO4 values without significant changes in δ 34 S SO4 [START_REF] Fontes | Oxygen, Carbon, Sulfur and Hydrogen stable isotopes in carbonate and sulfate mineral phases of Neogene evaporites, sediments and in interstitial waters[END_REF][START_REF] Pierre | Contribution ~ l'Jtude sgdimentologique et isotopique des evaporites messiniennes de la Mediterrane; implications geodynamiques[END_REF][START_REF] Ricchiuto | Stable Isotopic investigation of Messinian sulfate samples from DSDP. LEG 42. Eastern Mediterranean Sea[END_REF], this hypothesis seems highly unlikely as the abiotic equilibration between sulfate and water oxygen take about 20 Myr at normal marine pH (Lloyd, 1968;[START_REF] Longinelli | Oxygen-18 variations in sulfate ions in sea water and saline lakes[END_REF][START_REF] Turchyn | Oxygen isotopic composition of sulfate in deep sea pore fluid: evidence for rapid sulfur cycling[END_REF]. Moreover, the microbial sulfate reduction process is supported by the presence of pyrite at Sites 132, 654A and 968 (Pierre, 1982;Pierre and Rouchy, 1990;Blanc-Valleron et al., 1998) and the existence of filaments of possible microbial origin at Site 654A (Pierre and Rouchy, 1990).

The sulfate isotopic values reported by Longinelli (1979) and Pierre (1982) from the Upper Gypsum of Eraclea Minoa (Caltanissetta Basin, Sicily) are considerably more scattered than those from a recent study by García-Veigas et al., 2018;Fig. 10b). Such discrepancies are probably a consequence of different sample selection: García-Veigas et al. ( 2018) analyzed only pristine whitish selenite and balatino samples, while Longinelli (1979) and Pierre (1982) analyzed all types of gypsumbearing samples such as "gypsiferous marl" and gypsum laminae intercalated in carbonate or diatomaceous intervals. These less pristine samples probably contain high quantities of 34 S-depleted solid sulfides or diagenetic gypsum formed by oxidation of sulfides (see [START_REF] Liu | Sulfate sources constrained by sulfur and oxygen isotopic compositions in the upper reaches of the Xijiang River, China[END_REF] for more details on this process) and are therefore unlikely to be representative of the primary gypsum facies. Once these data are excluded, the Eraclea Minoa sulfate values (δ 18 O SO4 from 12.4 to 14.6‰ and δ 34 S SO4 from 21.0 to 22.3‰) suggest a marine origin of the sulfate and stable redox conditions during gypsum deposition (Fig 9.b;García-Veigas et al., 2018). Interestingly, the Eraclea Minoa sulfate values are in compliance with the isotopic values (δ 18 O SO4 =15.2 to 16.8‰; δ 34 S SO4 = 20.4 to 21.9‰) measured by Pierre (1982) in the Polemi Basin (Cyprus).

Hydrogen isotopes on molecular biomarkers

From the point of view of the application of organic geochemistry proxies, the Miocene Mediterranean Basin received little attention so far, with biomarker-based proxies that have been mostly applied to (a limited number of) pre-MSC sequences [START_REF] Tzanova | Cooling Mediterranean Sea surface temperatures during the Late Miocene provide a climate context for evolutionary transitions in Africa and Eurasia[END_REF]Herbert et al., 2016;[START_REF] Mayser | Precession driven changes in terrestrial organic matter input to the Eastern Mediterranean leading up to the Messinian Salinity Crisis[END_REF]Natalicchio et al., 2017Natalicchio et al., , 2019;;Vasiliev et al., 2019) and pre-Stage 3 sedimentary records (Lower Evaporites on Sicily, [START_REF] Andersen | Large and rapid climate variability during the Messinian salinity crisis: evidence from deuterium concentrations of individual biomarkers[END_REF]Vena del Gesso Basin, Sinninghe Damsté et al., 1995 andVasiliev et al., 2017;Levant Basin, Meilijson et al., 2019). To date, only one study analyzed Stage 3 samples (Vasiliev et al., 2017). This study used compound specific hydrogen isotope (δD) analyses, measured on both terrestrial (long chain C 29 and C 31 n-alkanes; [START_REF] Sachse | δD values of individual n-alkanes from terrestrial plants along a climatic gradient -implications for the sedimentary biomarker record[END_REF] and aquatic (alkenones; [START_REF] Englebrecht | Determination of sediment provenance at drift sites using hydrogen isotopes and unsaturation ratios in alkenones[END_REF] biomarkers from the gypsum beds of the Upper Gypsum at Eraclea Minoa to reconstruct the hydrological cycle during gypsum precipitation.

Both δD C29n-alkane and δD alkenones results (Fig. 10c) suggested that conditions in Sicily were significantly dryer than today, with highly enriched values of δD C29n-alkanes (up to -125‰). The δD alkenones varied between values suggesting evaporative conditions (-125‰) and values typical for present-day δD alkenones in the Mediterranean (-203‰) (Vasiliev et al., 2017).

No time-equivalent biomarker data from the open ocean settings are currently available. Instead, Vasiliev et al. (2017) compared their Mediterranean data with data from the Black Sea (DSDP 42B Hole and Taman peninsula; [START_REF] Vasiliev | Impact of the Messinian Salinity Crisis on Black Sea hydrology -insights from hydrogen isotopes on molecular biomarkers[END_REF][START_REF] Vasiliev | Recurrent phases of drought in the upper Miocene of the Black Sea region[END_REF]. The Upper Gypsum δD n-alkanes were more enriched when compared to their time equivalent deposits of the DSDP 42B 380 borehole of the Black Sea (-180‰). This probably reflects the more intracontinental position of the Black Sea which commonly translates into more depleted values for δD precipitation used by the vegetation, resulting in more depleted δD C29n -alkanes . However, there is a 30 to 40‰ enrichment relative to present in the δD n- alkanes (i.e. δD precipitation ) in both Mediterranean and Paratethys domains, indicating concurrent changes in both areas during the latest phase of the MSC.

Both the Mediterranean and Paratethyan samples contain δD alkenones with low values (~-200‰) (Fig. 10c) leading Vasiliev et al. (2017) to suggest that either the surface water from the Upper Gypsum was derived from the Black Sea, or that the Mediterranean and Paratethys were exchanging surface water during gypsum precipitation. Similarity between the relative contribution of the C 37 , C 38 and C 39 alkenones at Eraclea Minoa and one of the Black Sea samples may suggest common alkenone producers for the two areas, again supporting the idea of a Mediterranean-Paratethys connection during Stage 3 (Vasiliev et al., 2017).

A final speculative insight from this biomarker dataset is that the relative contribution of alkenones found in the Upper Gypsum of Eraclea Minoa is strikingly similar to present-day open marine samples, even though Emiliania huxleyi, the principal ocean alkenone producer today, did not exist in the late Miocene. Vasiliev et al. (2017) suggested that this could imply the existence of a connection to the open ocean during Upper Gypsum deposition in Sicily (i.e. throughout Stage 3; Fig. 3a).

Oxygen isotopes

Oxygen stable isotope data (δ 18 O) are available from bulk samples [START_REF] Rouchy | Paleoenvironmental changes at the Messinian-Pliocene boundary in the eastern Mediterranean (southern Cyprus basins): significance of the Messinian Lago-Mare[END_REF][START_REF] Rouchy | Late Messinian to Early Pliocene paleoenvironmental changes in the Melilla Basin (NE Morocco) and their relation to Mediterranean evolution[END_REF][START_REF] Rouchy | The end of the Messinian salinity crisis: evidences from the Chelif Basin (Algeria)[END_REF][START_REF] Pierre | Reconstruction of the paleoenvironmental changes around the Miocene-Pliocene boundary along a West-East transect across the Mediterranean[END_REF][START_REF] Cosentino | Orbitally forced paleoenvironmental and paleoclimate changes in the late postevaporitic Messinian of the central Mediterranean Basin[END_REF], gypsum [START_REF] Pierre | Isotope Composition of Messinian sediments from the Mediterranean Sea as indicators of paleoenvironments and diagenesis[END_REF][START_REF] Ricchiuto | Stable Isotopic investigation of Messinian sulfate samples from DSDP. LEG 42. Eastern Mediterranean Sea[END_REF]Lugli et al., 2007), mollusk shells [START_REF] Carnevale | Did the Mediterranean marine reflooding precede the Miocene/Pliocene boundary? Paleontological and geochemical evidence from upper Messinian sequences of Tuscany, Italy. Paleogeogr[END_REF][START_REF] Grunert | Estuarine Lago Mare fauna from the Tertiary Piedmont Basin indicates episodic Atlantic/Mediterranean exchange during the final stage of the Mediterranean Salinity Crisis[END_REF] and ostracod valves [START_REF] Cosentino | Orbitally forced paleoenvironmental and paleoclimate changes in the late postevaporitic Messinian of the central Mediterranean Basin[END_REF]Grossi et al., 2015).

For all the sub-basins for which there is latest Messinian data (e.g. Sites 974 and 975; Eraclea Minoa section, Sicily; Aghios Stefanos section, Corfu; Kalamaki section, Zakynthos; Pissouri Basin, Cyprus; [START_REF] Rouchy | Paleoenvironmental changes at the Messinian-Pliocene boundary in the eastern Mediterranean (southern Cyprus basins): significance of the Messinian Lago-Mare[END_REF][START_REF] Pierre | Reconstruction of the paleoenvironmental changes around the Miocene-Pliocene boundary along a West-East transect across the Mediterranean[END_REF], each has its own range of oxygen isotopic compositions and its own degree of variability. Values from above the Miocene/Pliocene boundary regain seawater values of 0.3 to ‰ (e.g. [START_REF] Pierre | Reconstruction of the paleoenvironmental changes around the Miocene-Pliocene boundary along a West-East transect across the Mediterranean[END_REF].

In marginal marine settings and lakes, the controls over δ 18 O are poorly constrained as oxygen does not respond simply to the freshwater flux, but to a combination of variables such as temperature, rainfall and evaporation (e.g. [START_REF] Placzek | Isotopic tracers of paleohydrologic change in large lakes of the Bolivian Altiplano[END_REF]. Freshwater input may contribute to the signal, resulting in δ 18 O more negative than seawater (0.3‰ to 0.8‰ SMOW; [START_REF] Dettman | The use of oxygen isotope variation in shells of estuarine mollusks as a quantitative record of seasonal and annual Colorado River discharge[END_REF], but under prevailing evaporating conditions it is likely that the δ 18 O will be primarily influenced by evaporation, leading to δ 18 O more positive than seawater (e.g. [START_REF] Dettman | The use of oxygen isotope variation in shells of estuarine mollusks as a quantitative record of seasonal and annual Colorado River discharge[END_REF], making any data very difficult to interpret. Furthermore, the lack of a unique δ 18 O signature for each water source makes oxygen isotopes a difficult tracer proxy to use.

APPENDIX D

Summary of the Stage 3 geochemical dataset

The variety of paleoenvironmental and connectivity proxies applied to MSC Stage 3 record provide valuable insights into the hydrological conditions during Stage 3. The more outstanding results from all discussed proxies are that: 1) Paleodepositional subaqueous environments where gypsum was precipitating and ostracods and biomarker-producers were thriving were strongly dominated by non-oceanic inputs; 2) an indisputable marine signal is absent and only regained above the M/P boundary.

Sulfate and oxygen isotopes are currently difficult to use for water provenance reconstruction because the non-marine sources (local and major rivers and Eastern Paratethys) that are likely to be of influence lack distinctive isotopic signatures and, especially for oxygen, respond to a combination of controls (e.g. temperature, rainfall, evaporation) with local variability. δ 34 S SO4 are claimed by several authors to be an evidence of the presence of an Atlantic inflow (δ 34 S SO4 =22‰; Turchyn and Schrag, 2004) in a Mediterranean strongly affected by non-marine waters [START_REF] Manzi | A new facies model for the Upper Gypsum of Sicily (Italy): chronological and paleoenvironmental constraits for the Messinian salinity crisis in the Mediterranean[END_REF](Manzi et al., , 2016a;;García-Veigas et al., 2018 among others). However, the same values can be obtained by means of the recycling of PLG deposits (~23‰; Lu et al., 2001;Lugli et al., 2010;García-Veigas et al., 2018).

Similarities between the δD alkenones of the Upper Gypsum at Eraclea Minoa and coeval Black Sea sediments and δD n-alkanes similar to presentday marine settings, suggest that Eastern Paratethys and the Atlantic were simultaneously contributing to the Mediterranean hydrological budget. 87 Sr/ 86 Sr isotope ratios are a useful water-mass tracer because each water body carries a unique Sr isotope fingerprint (see subsection 8.1.1). Our plotting of Stage 3 87 Sr/ 86 Sr isotope values (Fig. 10a) highlights the large geographical variability of the values and the sharp division between Sr isotope ratios measured in marginal basins versus those in intermediate-deep water locations. This is only noticeable in substage 3.2, since no (or not enough) material suitable for Sr analysis is present in substage 3.1 deposits from the marginal basins. Some authors see this variability as an indication of isolated subbasins with unique hydrological conditions driven by their catchment rivers (e.g. Müller et al., 1990;[START_REF] Müller | Origin and age of the Mediterranean Messinian evaporites: implications from Sr isotopes[END_REF]Ryan, 2009). If some degree of connection was present, it involved only neighbouring basins (e.g. Tyrrhenian subbasins; Müller et al., 1990;[START_REF] Müller | Origin and age of the Mediterranean Messinian evaporites: implications from Sr isotopes[END_REF]. A recent comparison of the Sr isotope record of the Spanish marginal basins of Sorbas, Nijar and Vera with the Sr isotope ratios likely to have typified the local riverine sources demonstrated that a local sourcesmixed signal expected from an endorheic lake in that location is absent. In this instance mixing of intrabasinal water sources with a nonmarine Mediterranean water mass is used to explain the measured values (Andreetto et al., 2021). If this explanation is more widely applicable, then it may result in a re-interpretation of the spread of Sr isotope data from the latest Messinian interval.

To conclude, geochemical proxies have great potential to test the different scenarios, but data are currently too numerically and geographically limited to be robust.

Paleoenvironmental scenarios for freshening the salt giant: desiccated versus full Mediterranean

The riddle of the Mediterranean environmental and hydrological conditions during Stage 3 is a highly debated topic and it is key to understanding the means by which open marine conditions were restored at the base of the Zanclean and on the potential impact that the Atlantic-Mediterranean re-connection had on the Atlantic and global climate (Flecker et al., 2015;[START_REF] Capella | Mediterranean isolation preconditioning the Earth System for late Miocene climate cooling[END_REF]. In this chapter, the paleoenvironmental scenarios, in terms of base-level position (desiccated or full Mediterranean) and hydrological configuration (connections to the Atlantic and/or Paratethys), proposed for the Mediterranean during Stage 3 are described, as well as the different timings of the reflooding (instantaneous, gradual, step-like increments). The low-salinity Stage followed the hypersaline Stage 2 and the transition between the two likely influences the plausibility of the various paleoenvironmental scenarios proposed for the terminal stage. We therefore first summarize the current understanding of the configuration of the Mediterranean during Stage 2 and the enduring controversies (see Roveri et al., 2014a for a more extensive review).

7. : formation of the Mediterranean salt giant Numerical modelling based on hydrological budget calculations shows that in order to reach salinity levels compatible with halite saturation and to accumulate the substantial thicknesses of halite observed in the seismic profiles (Ryan, 1973;Haq et al., 2020), the Atlantic-Mediterranean gateway needs to have permitted inflow from the Atlantic, but may have completely blocked outflow [START_REF] Blanc | The opening of the Plio-Quaternary Gibraltar Strait: assessing the size of a cataclysm[END_REF]Krijgsman and Meijer, 2008). Numerical models also showed that without Atlantic inflow into the Mediterranean Sea its base level is forced to drop on time scales in the order of a few thousand years by virtue of the basin's negative hydrological budget, where more water is lost to the atmosphere by evaporation than is received from rainfall and river runoff (e.g. [START_REF] Meijer | A quantitative analysis of the desiccation and re-filling of the Mediterranean during the Messinian Salinity Crisis[END_REF]Krijgsman and Meijer, 2008;Simon et al., 2017). The idea of a drawdown is supported by several arguments: (1) the widespread presence, from the margins to the slopes, of the Messinian Erosional Surface cutting through Stage 1 and pre-MSC deposits and canyon incisions following today's drainage networks (e.g. [START_REF] Chumakov | Pliocene and Pleistocene deposits of the Nile valley in Nubia and upper Egypt[END_REF][START_REF] Clauzon | Le canyon messinien du Rhoˆne: Une preuve decisive du "dessicated deep basin model[END_REF][START_REF] Lofi | Erosional processes and paleo-environmental changes in the Western Gulf of Lions (SW France) during the Messinian Salinity Crisis[END_REF]Lofi et al., , 2011aLofi et al., , 2011b;;[START_REF] Loget | Mesoscale fluvial erosion parameters deduced from modelling the Mediterranean sea-level drop during the Messinian (late Miocene)[END_REF]Maillard et al., 2006[START_REF] Maillard | New onshore/ offshore evidence of the Messinian Erosion Surface from key areas: The Ibiza-Balearic Promontory and the Orosei-Eastern Sardinian margin. Découverte de la surface d'érosion messinienne onshore/offshore dans deux lieux clés: le Promontoire Baléares (Ibiza) et la marge est-sarde (Orosei)[END_REF][START_REF] Andreetto | Impact of pulsed Atlantic water inflow into the Alboran Basin at the time of the Zanclean flooding[END_REF][START_REF] Just | Erosion of continental margins in the Western Mediterranean due to sea-level stagnancy during the Messinian Salinity Crisis[END_REF][START_REF] Urgeles | New constraints on the Messinian sealevel drawdown from 3D seismic data of the Ebro margin, western Mediterranean[END_REF][START_REF] Amadori | Restored topography of the Po Plain-Northern Adriatic region during the Messinian base level drop-Implications for the physiography and compartmentalization of the paleo-Mediterranean basin[END_REF][START_REF] Lymer | The Western Tyrrhenian Sea revisited: New evidence for a rifted basin during the Messinian Salinity Crisis[END_REF][START_REF] Cazzini | New seismic evidence of the Messinian paleomorphology beneath Lake Maggiore area (Italy)[END_REF]Figs. 5e, 7e); (2) their morphology interpreted as subaerial in origin; [START_REF] Lein | Isotopic composition of sulfur compounds in sediments of the Pacific[END_REF] the clastic fans at the outlet of the valleys onlapped by Stage 3 deposits and interpreted as fluvial accumulations (e.g. [START_REF] Lofi | Erosional processes and paleo-environmental changes in the Western Gulf of Lions (SW France) during the Messinian Salinity Crisis[END_REF]Maillard et al., 2006;[START_REF] Pellen | The Messinian Ebro River incision[END_REF]. A number of studies have tried to quantify the magnitude of the sea-level fall by compensating for the isostatic vertical motion since the Messinian to obtain the original depth of the erosional features and Messinian deposits. However, this depends on the assumptions about when the drawdown occurred relative to the halite precipitation: before (e.g. [START_REF] Cartwright | Initiation of gravitational collapse of an evaporitic basin margin: the Messinian saline giant, Levant Basin, eastern Mediterranean[END_REF][START_REF] Bache | Messinian erosional and salinity crisis: view from the Provence basin (Gulf of Lions, western Mediterranean)[END_REF][START_REF] Bache | A twostep process for the reflooding of the Mediterranean after the Messinian Salinity Crisis[END_REF], during (e. g. [START_REF] Ryan | Modeling the magnitude and timing of evaporative drawdown during the Messinian salinity crisis[END_REF]Ryan, , 2009) ) or after (e.g. [START_REF] Ryan | Messinian badlands on the southeastern margin of the Mediterranean Sea[END_REF][START_REF] Bertoni | Major erosion at the end of the Messinian Salinity Crisis: evidence from the Levant Basin, Eastern Mediterranean[END_REF]Lofi et al., 2011aLofi et al., , 2011b)). How shallow the Mediterranean became during Stage 2 is also a matter of disagreement. Estimates in the Western Mediterranean vary from a maximum drawdown of 2500 m [START_REF] Ryan | Quantitative evaluation of the depth of the western Mediterranean before, during and after the late Miocene salinity crisis[END_REF] to 1000-1500 m [START_REF] Bache | A twostep process for the reflooding of the Mediterranean after the Messinian Salinity Crisis[END_REF] in the Gulf of Lion, 800-1200 m in the Balearic promontory (Mas et al., 2018b) and 400 m in the Ebro delta region [START_REF] Frey-Martinez | 3D seismic interpretation of the Messinian Unconformity in the Valencia Basin, Spain[END_REF]. A later backstripping analysis of this delta yielded a drawdown of ~1300 m [START_REF] Urgeles | New constraints on the Messinian sealevel drawdown from 3D seismic data of the Ebro margin, western Mediterranean[END_REF]. East of the Sicily sill, backstripping studies estimated base-level drops of 1800-2000 m in the Ionian basin [START_REF] Micallef | Evidence of the Zanclean megaflood in the eastern Mediterranean Basin[END_REF][START_REF] Micallef | Geomorphic evolution of the Malta Escarpment and implications for the Messinian evaporative drawdown in the eastern Mediterranean Sea[END_REF][START_REF] Camerlenghi | Seismic markers of the Messinian salinity crisis in the deep Ionian Basin[END_REF][START_REF] Spatola | A single-stage megaflood at the termination of the Messinian salinity crisis: geophysical and modelling evidence from the eastern Mediterranean Basin[END_REF], 800-900 m in the Adriatic foredeep and Po plain [START_REF] Ghielmi | Late Miocene-Middle Pleistocene sequences in the Po Plain-Northern Adriatic Sea (Italy): the stratigraphic record of modification phases affecting a complex foreland basin[END_REF][START_REF] Amadori | Restored topography of the Po Plain-Northern Adriatic region during the Messinian base level drop-Implications for the physiography and compartmentalization of the paleo-Mediterranean basin[END_REF]), 800-1300m (Ben-Gal et al., 2005), 600 [START_REF] Druckman | The buried Afiq Canyon (eastern Mediterranean, Israel): a case study of a Tertiary submarine canyon exposed in Late Messinian times[END_REF] and m [START_REF] Cartwright | Initiation of gravitational collapse of an evaporitic basin margin: the Messinian saline giant, Levant Basin, eastern Mediterranean[END_REF] in the Levant Basin.

None of these quantifications could unequivocally constrain the timing of the drawdown within the MSC sequence, but numerical modeling studies show that, if the blocking of the outflow was controlled by a tectonic uplift counteracted by inflow erosion across the Strait of Gibraltar, then the expected drawdown of the Mediterranean Sea should be moderate (< 400 m; and possibly harmonic) due to an equilibrium between incision and uplift before the complete blocking of inflow and larger (up to complete desiccation) only after tectonic uplift overcame incision rates (Garcia-Castellanos and Villaseñor, 2011). The same model suggests that the initiation of halite precipitation might overlap in time with the late primary gypsum deposition, right before the full

APPENDIX D

disconnection from the Atlantic Ocean.

The interpretation of the deep evaporites and their associated seismic markers (erosional surfaces and deep engravings along the shelf-slope systems) is not straightforward. Recently, it was suggested that the deep evaporitic facies and the seismic morphological features could have been produced without a significant drop of the Mediterranean base-level, therefore promoting the persistence of a relatively deepwater Mediterranean basin even during halite deposition [START_REF] Lugli | Evidence of clastic evaporites in the canyons of the Levant Basin (Israel): implications for the Messinian Salinity Crisis[END_REF][START_REF] Lugli | The deep record of the Messinian salinity crisis: evidence of a non-desiccated Mediterranean Sea[END_REF]Roveri et al., 2014b). For example, Roveri et al. (2014c) proposed that downslope flows of dense, hypersaline waters sourced from evaporation in shallower water areas could have generated both the observed shelf-slope erosion and have created a deep brine, supersaturated in the ions necessary for precipitating halite. These subaqueous hyperpycnal flows are consistent with the observed clastic evaporites that filled the Levant margin canyons [START_REF] Lugli | Evidence of clastic evaporites in the canyons of the Levant Basin (Israel): implications for the Messinian Salinity Crisis[END_REF] and, more generally, with the widespread presence of Complex Units at the outlet of the MES drainage systems (see [START_REF] Lofi | Erosional processes and paleo-environmental changes in the Western Gulf of Lions (SW France) during the Messinian Salinity Crisis[END_REF]Lofi et al., , 2011aLofi et al., , 2011b;;Lofi, 2018). These sediments are dominated by reworked PLG that would have been exposed by a sea-level fall as little as 200 m (Lugli et al., 2010). However, the hypersaline environment that is presumed to be established by these hyperpycnal flows during the deposition of the RLG is in contrast with the occurrence of the Paratethyan ostracod L. muelleri within the clastic evaporites (RLG) in several marginal sections (e.g. Adana Basin, [START_REF] Faranda | Messinian paleoenvironmental changes in the easternmost Mediterranean Basin: Adana Basin, southern Turkey[END_REF][START_REF] Radeff | Stratigraphic architecture of the upper Messinian deposits of the Adana Basin (southern Turkey): implications for the Messinian salinity crisis and the Taurus petroleum system[END_REF][START_REF] Radeff | Sedimentary evidence for late Messinian uplift of the SE margin of the Central Anatolian Plateau: Adana Basin, southern Turkey[END_REF].

Whatever the state of Mediterranean base-level during Stage 2, the more commonly used chronostratigraphic model for the MSC (Fig. 1a; Roveri et al., 2014a) states that massive halite precipitation ceased at 5.55 Ma and was superseded by an environment that, with precession periodicity (Fig. 3a), cycled between gypsum precipitation and conditions that saw fresh-brackish organisms thriving. The question is whether these conditions cycled homogeneously in several isolated lakes or in basins largely connected to the same Atlantic and Eastern Paratethys-influenced water mass (Fig. 11).

7.2. : resumption of (upper) gypsum precipitation and Paratethys fauna invasion

An isolated Mediterranean dotted by sabkhas and lakes

The first and long-lasting paleoenvironmental interpretation of the evaporite-bearing UG/UU units and (possibly) time-equivalent evaporite-free units (e.g. LM Unit in Malaga, Sorbas and Zorreras Mb. in Sorbas, Feos Fm. in Nijar, Cassano Spinola Conglomerates in Piedmont, San Donato/Colombacci fms. in the Apennines, Handere Fm. in Turkey) envisaged their sedimentation in a Mediterranean mostly isolated from the Paratethys (which may have added water only to some basins in the Eastern Mediterranean) and totally isolated from the Atlantic where, in each subbasin, continental settings (e.g. alluvial plains, river channels, alluvial fans, playa lakes, sabkhas) alternated/interfingered with shallow, endorheic lakes (Figs. 11a,c; e.g. [START_REF] Ruggieri | La serie marine pliocenica e quaternaria della Val Marecchia: Atti Accad[END_REF]Ruggieri, , 1967;;[START_REF] Decima | Comments on late Messinian microfaunas in several sections from Sicily[END_REF][START_REF] Decima | Late Miocene evaporites of the Central Sicilian Basin[END_REF][START_REF] Friedman | Petrographic data and comments on the depositional environment of the Miocene sulfates and dolomites at Sites 124, 132, and 134, western Mediterranean Sea[END_REF]Hsü et al., 1973a, 1973b, 1973c, Hsü et al., 1978a[START_REF] Hsü | Initial reports of the deep sea drilling project: DSDP volume XLII Part[END_REF]Ryan et al., 1973;[START_REF] Selli | An outline of the Italian Messinian[END_REF][START_REF] Sturani | A fossil eel (Anguilla sp.) from the Messinian of Alba (Tertiary Piedmont Basin). Paleoenvironmental and paleogeographic implications[END_REF][START_REF] Sissingh | Aspects of late Cenozoic evolution of the South Aegean ostracode fauna. Paleogeogr., Paleoclimatol[END_REF][START_REF] Benson | The paleoecology of the ostracods of DSDP Leg 42A[END_REF][START_REF] Bossio | Some aspects of the upper Miocene in Tuscany[END_REF]Cita et al., 1978[START_REF] Cita | Messinian Paleoenvironments: New Evidence from the Tyrrhenian Sea (ODP LEG 107)[END_REF][START_REF] Ricchiuto | Stable Isotopic investigation of Messinian sulfate samples from DSDP. LEG 42. Eastern Mediterranean Sea[END_REF][START_REF] Ryan | Messinian badlands on the southeastern margin of the Mediterranean Sea[END_REF][START_REF] Ryan | Modeling the magnitude and timing of evaporative drawdown during the Messinian salinity crisis[END_REF]Ryan, , 2009;;[START_REF] Cita | Sedimentation in the latest Messinian at Capo Rossello (Sicily)[END_REF][START_REF] Orszag-Sperber | Le Miocène terminal et le Pliocène inférieur au sud de Chypre, livret-guide, 5e séminaire sur le Messinien[END_REF][START_REF] Ghibaudo | Litostratigrafia, sedimentologia ed evoluzione tettonico-sedimentaria dei depositi miocenici del margine sud-orientale del bacino terziario ligure-piemontese (Valli Borbera, Scrivia e Lemme). In memoria di Carlo Sturani[END_REF]Müller et al., 1990;Benson and Rakic-El Bied, 1991;Benson et al., 1991;[START_REF] Müller | Origin and age of the Mediterranean Messinian evaporites: implications from Sr isotopes[END_REF][START_REF] Orszag-Sperber | La transition Messinien-Pliocène en Méditerranée orientale (Chypre): la période du Lago-Mare et sa signification[END_REF][START_REF] Rouchy | Paleoenvironmental changes at the Messinian-Pliocene boundary in the eastern Mediterranean (southern Cyprus basins): significance of the Messinian Lago-Mare[END_REF][START_REF] Rouchy | Late Messinian to Early Pliocene paleoenvironmental changes in the Melilla Basin (NE Morocco) and their relation to Mediterranean evolution[END_REF][START_REF] Rouchy | The end of the Messinian salinity crisis: evidences from the Chelif Basin (Algeria)[END_REF][START_REF] Blanc | The opening of the Plio-Quaternary Gibraltar Strait: assessing the size of a cataclysm[END_REF][START_REF] Lofi | Erosional processes and paleo-environmental changes in the Western Gulf of Lions (SW France) during the Messinian Salinity Crisis[END_REF], Lofi et al., 2011b;[START_REF] Bassetti | Evolution of depositional environments after the end of Messinian Salinity Crisis in Nijar basin (SE Betic Cordillera)[END_REF]Rouchy and Caruso, 2006;[START_REF] Bertoni | Major erosion at the end of the Messinian Salinity Crisis: evidence from the Levant Basin, Eastern Mediterranean[END_REF][START_REF] Cameselle | Large-scale margin collapse during Messinian early sea-level drawdown: the SW Valencia trough, NW Mediterranean[END_REF][START_REF] Amadori | Restored topography of the Po Plain-Northern Adriatic region during the Messinian base level drop-Implications for the physiography and compartmentalization of the paleo-Mediterranean basin[END_REF][START_REF] Camerlenghi | Seismic markers of the Messinian salinity crisis in the deep Ionian Basin[END_REF][START_REF] Kartveit | Evidence of sea level drawdown at the end of the Messinian salinity crisis and seismic investigation of the Nahr Menashe unit in the northern Levant Basin, offshore Lebanon[END_REF][START_REF] Madof | Discovery of vast fluvial deposits provides evidence for drawdown during the late Miocene Messinian salinity crisis[END_REF][START_REF] Ben Moshe | Estimating drawdown magnitudes of the Mediterranean Sea in the Levant basin during the Lago Mare stage of the Messinian Salinity Crisis[END_REF][START_REF] Caruso | The late Messinian "Lago-Mare" event and the Zanclean Reflooding in the Mediterranean Sea: new insights from the Cuevas del Almanzora section (Vera Basin, South-Eastern Spain)[END_REF][START_REF] Cazzini | New seismic evidence of the Messinian paleomorphology beneath Lake Maggiore area (Italy)[END_REF]Raad et al., 2021). The full disconnection is also supported by observations that support an abrupt Zanclean reflooding (e.g. [START_REF] Blanc | The opening of the Plio-Quaternary Gibraltar Strait: assessing the size of a cataclysm[END_REF][START_REF] Micallef | Evidence of the Zanclean megaflood in the eastern Mediterranean Basin[END_REF][START_REF] Micallef | Geomorphic evolution of the Malta Escarpment and implications for the Messinian evaporative drawdown in the eastern Mediterranean Sea[END_REF][START_REF] Garcia-Castellanos | The Zanclean megaflood of the Mediterranean-Searching for independent evidence[END_REF][START_REF] Spatola | A single-stage megaflood at the termination of the Messinian salinity crisis: geophysical and modelling evidence from the eastern Mediterranean Basin[END_REF], since a rapid outburst flood requires a large sea level difference prior to the flood that can only be developed in a scenario of a full Mediterranean-Atlantic disconnection [START_REF] Garcia-Castellanos | Catastrophic flood of the Mediterranean after the Messinian salinity crisis[END_REF]Garcia-Castellanos and Villaseñor, 2011). Although rarely explicitly stated, all these studies must assume that: 1) all Paratethyan biota (and possibly other organisms of undisclosed provenance like diatoms) migrated passively via aquatic migratory birds across the entire Mediterranean (Fig. 11a; [START_REF] Benson | The paleoecology of the ostracods of DSDP Leg 42A[END_REF]Benson and Rakic-El Bied, 1991;[START_REF] Caruso | The late Messinian "Lago-Mare" event and the Zanclean Reflooding in the Mediterranean Sea: new insights from the Cuevas del Almanzora section (Vera Basin, South-Eastern Spain)[END_REF]; 2) chemical and physical conditions (brackish water and water depth not exceeding 100 m; e.g. Hajós, 1973;[START_REF] Gliozzi | Late Messinian lago-mare ostracod paleoecology: a correspondence analysis aroach[END_REF]) that allowed alternated conditions suitable for gypsum to precipitate and Paratethyan biota and euryhaline benthic foraminifera to thrive were related to changes in the local freshwater budget; 3) The marine isotopic signals in UU/UG gypsum (Fig. 10) are entirely the reflection of the lithologies that are leached by continental waters in surficial and/or underground drainage areas (e.g. Ryan, 2009;Raad et al., 2021); 4) Stage 3 gypsum precipitated in extremely shallow-water (playa lakes) to completely dried environments (sabkhas) and the excessive sulfate necessary is completely derived from "clastic reworking, dissolution, re-precipitation and diagenesis of materials belonging to the PLG and halite of the previous MSC Stage 2" (Ryan, 2009).

Observations supporting a Mediterranean isolated throughout Stage 3 and only at the mercy of local freshwater inputs (Fig. 11a) are: (1) the lack of evidence for in situ marine fauna and flora in UU (e.g. Ryan et al., 1973;Hsü et al., 1978a;[START_REF] Cita | Messinian Paleoenvironments: New Evidence from the Tyrrhenian Sea (ODP LEG 107)[END_REF]Ryan, 2009;Lofi et al., 2011a);

(2) the shallow-water mode of life and highly likely in-situ nature of ostracods and euryhaline, shallow-water benthic foraminifera observed in DSDP/ODP wells from intermediate and deep basins (e.g. Cita et al., 1978;[START_REF] Iaccarino | Paleoenvironment of uppermost Messinian sequences in the western Mediterranean (Sites 974, 975, and 978[END_REF]Figs. 9a -c); (3) the bathymetric contrast (up to several hundred meters) between the late Messinian paleoenvironments and the marine Zanclean on top (e.g. [START_REF] Cita | Sedimentation in the latest Messinian at Capo Rossello (Sicily)[END_REF][START_REF] Bonaduce | Paleoecological interpretation of the latest Messinian sediments from southern Sicily (Italy)[END_REF][START_REF] Caruso | The late Messinian "Lago-Mare" event and the Zanclean Reflooding in the Mediterranean Sea: new insights from the Cuevas del Almanzora section (Vera Basin, South-Eastern Spain)[END_REF]; (4) the presence of paleosols in Cyprus [START_REF] Orszag-Sperber | La transition Messinien-Pliocène en Méditerranée orientale (Chypre): la période du Lago-Mare et sa signification[END_REF][START_REF] Rouchy | Paleoenvironmental changes at the Messinian-Pliocene boundary in the eastern Mediterranean (southern Cyprus basins): significance of the Messinian Lago-Mare[END_REF] and on the crest of the Eratosthenes seamount (Robertson, 1998a(Robertson, , 1998b)); ( 5) the erosional features preserved both offshore on the continental shelves and lower-middle slope domain and interpreted in most seismic stratigraphic studies as the result of subaerial exposure (e. g. [START_REF] Lofi | Erosional processes and paleo-environmental changes in the Western Gulf of Lions (SW France) during the Messinian Salinity Crisis[END_REF]Lofi et al., 2011b;[START_REF] Lymer | The Western Tyrrhenian Sea revisited: New evidence for a rifted basin during the Messinian Salinity Crisis[END_REF][START_REF] Ben Moshe | Estimating drawdown magnitudes of the Mediterranean Sea in the Levant basin during the Lago Mare stage of the Messinian Salinity Crisis[END_REF]; (6) the pinching out of the UU/BU units towards evaporitefree pre-Messinian structural highs (Figs. 7b-g;Figs. 8a, e;Ryan, 2009;[START_REF] Lymer | The Western Tyrrhenian Sea revisited: New evidence for a rifted basin during the Messinian Salinity Crisis[END_REF][START_REF] Camerlenghi | Seismic markers of the Messinian salinity crisis in the deep Ionian Basin[END_REF]Raad et al., 2021); (7) the more abundant terrigenous clasts and reworked calcareous fossils in Stage 3 samples compared to the overlying, deep-water Pliocene (Ryan et al., 1973;[START_REF] Hsü | Initial reports of the deep sea drilling project: DSDP volume XLII Part[END_REF]Ryan, 2009); (8) the erosional nature of the M-reflector/TES/IMTS in the Levant Basin (Figs. 8e-g), by some linked to subaerial exposure of the Levant seafloor (e.g. [START_REF] Bertoni | Major erosion at the end of the Messinian Salinity Crisis: evidence from the Levant Basin, Eastern Mediterranean[END_REF]Lofi et al., 2011aLofi et al., , 2011b;;Maillard et al., 2011a) before the emplacement of deposits interpreted as fluvial from seismic observations [START_REF] Bowman | A comprehensive review of the MSC facies and their origins in the offshore Sirt Basin, Libya[END_REF][START_REF] Radeff | Sedimentary evidence for late Messinian uplift of the SE margin of the Central Anatolian Plateau: Adana Basin, southern Turkey[END_REF][START_REF] Leila | Depositional facies controls on the diagenesis and reservoir quality of the Messinian Qawasim and Abu Madi formations, onshore Nile Delta, Egypt[END_REF][START_REF] Kartveit | Evidence of sea level drawdown at the end of the Messinian salinity crisis and seismic investigation of the Nahr Menashe unit in the northern Levant Basin, offshore Lebanon[END_REF][START_REF] Madof | Discovery of vast fluvial deposits provides evidence for drawdown during the late Miocene Messinian salinity crisis[END_REF]. Furthermore, (9) isolated hydrological circuits with unique chemical composition are regarded by [START_REF] Camerlenghi | Seismic markers of the Messinian salinity crisis in the deep Ionian Basin[END_REF] as the most plausible explanation for the W-E change in the MSC sedimentary expression in the deep basins, represented by the trilogy LU-MU-UU in the Algero-Balearic and Liguro-Provençal basins, missing the LU in the Tyrrhenian and (possibly) Ionian basins, by terrigenous deposits with hiatuses in the WAB and Adriatic foredeep and by halite, anhydrite and clastics in the Levant Basin (Interbedded and Argillaceous evaporites of Meilijson et al., 2019;Fig. 3b).

The main problems with the isolated scenario lasting throughout Stage 3 are: (1) it does not provide an explanation neither for the homogeneity of Paratethyan ostracod assemblages in the marginal basins (e.g. [START_REF] Gliozzi | Paratethyan ostracod immigrants in Italy during the Late Miocene[END_REF]Stoica et al., 2016), an aspect difficult to explain when fauna migration takes place passively via either birds or wind, nor for the biomarkers (Vasiliev et al., 2017), which cannot be transported effectively by aquatic birds; (2) it does not explain the mismatch between 87 Sr/ 86 Sr isotope ratios measured on marginal ostracods and Sr values expected from endorheic lakes fed with local freshwaters (e.g. Andreetto et al., 2021); 3) it misses to substantiate, with geochemical arguments, the precipitation of gypsum in lakes, a process that is everything but straightforward (see Warren, 2016 for insights); 4) except for the salt-bearing basins, the source(s) of solutes which makes freshwater-fed endorheic lakes brackish and causes similar physico-chemical conditions to exist in each lake is also difficult to explain in the context of a Mediterranean only at the mercy of local rivers.

The half-full, density-stratified Mediterranean scenarios

An alternative concept to the isolated scenario envisages the Mediterranean connected with the Atlantic and/or the Eastern Paratethys and relatively full of water connecting the different subbasins (Fig. 11b). To our knowledge, this scenario was first developed by [START_REF] Mcculloch | Sr isotope constraints on the Mediterranean environment at the end of the Messinian salinity crisis[END_REF] on the basis of the similar 87 Sr/ 86 Sr ratios from marginal (Spain and Cyprus) and deep (Levantine and Algero-Balearic) basins. This intuition was a significant departure from the far more in vogue desiccated scenario (see conclusion of Hsü et al., 1973b), and for this was long overlooked. Sr isotope ratios lower than contemporary ocean water led [START_REF] Mcculloch | Sr isotope constraints on the Mediterranean environment at the end of the Messinian salinity crisis[END_REF] to conclude that a brackish water mass created by the mixing of water from the peri-Mediterranean rivers (e.g. Nile, Rhône and African rivers that no longer flow today, etc.; see [START_REF] Griffin | Aridity and humidity: two aspects of the late Miocene climate of North Africa and the Mediterranean[END_REF][START_REF] Gladstone | The Mediterranean hydrologic budget from a Late Miocene global climate simulation[END_REF] with water of the Eastern Paratethys filled the Mediterranean, resembling the Caspian Sea today. This conclusion is consistent with the impoverished (or absent) marine fauna and flora of Stage 3 sediments and the enhanced assemblage of fresh-brackish water biota (see subsection 5.7; Figs. 9a-c), but is problematic as a viable origin for Stage 3 gypsum to precipitate at depth. Furthermore, climate models for the late Miocene fail to fill the Mediterranean Basin with fluvial and Paratethys waters alone [START_REF] Gladstone | The Mediterranean hydrologic budget from a Late Miocene global climate simulation[END_REF]Marzocchi et al., 2016[START_REF] Marzocchi | Precessional drivers of late Miocene Mediterranean sedimentary sequences: African summer monsoon and Atlantic winter storm tracks[END_REF]Simon et al., 2017). A marine contribution is therefore required to fill the Mediterranean (Marzocchi et al., 2016). In the event, the contribution is most likely to have derived from the Atlantic via the Gibraltar Corridor (Flecker et al., 2015;Booth-Rea et al., 2018;Krijgsman et al., 2018) either through a karst system (Krijgsman et al., 2018) or an emerged volcanic archipelago in the Alborán Basin (Booth-Rea et al., 2018). In fact, although an Indian Ocean contribution was proposed (Cita et al., 1978;Hsü et al., 1978a) and the possibility discussed (Ryan, 2009;[START_REF] Vai | Over half a century of Messinian salinity crisis[END_REF], palinspastic reconstructions concluded that the Neo-Tethys Mediterranean-Indian Ocean connection via southern Turkey and Iran already closed before the Tortonian [START_REF] Rögl | Paleogeographic considerations for Mediterranean and Paratethys seaways (Oligocène to Miocene)[END_REF][START_REF] Popov | Lithological-paleogeographic maps of Paratethys-10 maps late Eocene to pliocene[END_REF][START_REF] Gargani | Evaporite accumulation during the Messinian Salinity Crisis: the Suez rift case[END_REF][START_REF] Bialik | Two-step closure of the Miocene Indian Ocean Gateway to the Mediterranean[END_REF][START_REF] Gülyüz | Paleomagnetic constraints on the early Miocene closure of the southern Neo-Tethys (Van region; East Anatolia): Inferences for the timing of Eurasia-Arabia collision[END_REF], while a seaway via the Red Sea and Gulf of Aden, although not completely ruled out (e.g. [START_REF] Schütz | Structure and stratigraphy of the Gulf of Suez, Egypt, in Interior Rift Basins[END_REF][START_REF] Bosworth | The Red Sea and Gulf of Aden Basins[END_REF][START_REF] Gargani | Evaporite accumulation during the Messinian Salinity Crisis: the Suez rift case[END_REF]Ryan, 2009), is highly contested (e.g. [START_REF] Meulenkamp | Tertiary paleogeography and tectonostratigraphic evolution of the Northern and Southern Peri-Tethys platforms and the intermediate domains of the African-Eurasian convergent plate boundary zone[END_REF][START_REF] Segev | Late Oligocene and Miocene different seaways to the Red Sea-Gulf of Suez rift and the Gulf of Aqaba-Dead Sea basins[END_REF].

In light of this, Roveri et al. (2014c), [START_REF] Gvirtzman | Intra-Messinian truncation surface in the Levant Basin explained by subaqueous dissolution[END_REF]), Vasiliev et al. (2017), García-Veigas et al. (2018) and Grothe et al. (2020) suggested that the Mediterranean was likely density-stratified during this interval as a result of the simultaneous influx of isotopicallydifferent marine and non-marine (major Mediterranean rivers and Eastern Paratethys) water sources (Fig. 11b). This connectivity framework resulted in a brackish layer carrying low-salinity (mostly Paratethyan) biota [START_REF] Gliozzi | Paratethyan ostracod immigrants in Italy during the Late Miocene[END_REF]Stoica et al., 2016;[START_REF] Andreetto | Migration of the dinoflagellate Galeacysta etrusca and its implications for the Messinian Salinity Crisis[END_REF]Grothe et al., , 2020;;Figs. 9a-b) to lay on top of a more saline layer formed by Atlantic-derived seawater from which UU/UG gypsum , that facies analyses demonstrated to result from subaqueous deposition (Hardie and Lowenstein, 2004;[START_REF] Lugli | The deep record of the Messinian salinity crisis: evidence of a non-desiccated Mediterranean Sea[END_REF], precipitated at intermediate and greater depths (e.g. García-Veigas et al., 2018). A dense, anoxic deep-water mass, possibly inherited from Stage 2, is envisaged at the bottom of the Mediterranean by Marzocchi et al. (2016) and García-Veigas et al. (2018), albeit without conclusive arguments, and by [START_REF] Gvirtzman | Intra-Messinian truncation surface in the Levant Basin explained by subaqueous dissolution[END_REF] following the observation that the tilted halite body of the Levant Basin was simultaneously eroded landward and preserved basinward (Fig. 8f).

This scenario accounts for the erosive/non-depositional features (Figs. 5e, 6a, e) and continental/lacustrine facies (Figs. 5a-b, d-g) widespread around the margins and shelves and suggestive of a Mediterranean base-level somewhat lower than the Atlantic level suggesting a one-way inflow from both the Atlantic and the Eastern Paratethys after Stage 2 (e.g. Marzocchi et al., 2016;Figs. 11b ,c), a connectivity configuration that effectively translates in a half-full Mediterranean (e.g. Krijgsman and Meijer, 2008). Refilling as a result of persistent Atlantic inflow, in part perhaps because of the latest Messinian deglaciation (see subsection 2.2; Van der Laan et al., 2006;[START_REF] Hilgen | Astronomical tuning as the basis for high resolution chronostratigraphy: the intricate history of the Messinian Salinity Crisis[END_REF], would have resulted in the establishment of two-way exchange first with the Paratethys at some point during the Lago-Mare phase and later, i.e. slightly before or at the Messinian/Zanclean boundary, with the Atlantic Ocean (Fig. 11c; Marzocchi et al., 2016). The moment the Mediterranean base-level reached the sill with the adjacent water body (Paratethys and Atlantic) and a two-way exchange was initiated, the density contrast will have prompted an enhanced inflow into the Mediterranean (Marzocchi et al., 2016). The overall transgressive trend leading to the Zanclean marine replenishment was accompanied by base-level fluctuations in the order of 400±100 m every precessional cycle (Fig. 11c; [START_REF] Fortuin | The Messinian of the Nijar basin (SE Spain): sedimentation, depositional environments and paleogeographic evolution[END_REF][START_REF] Ben Moshe | Estimating drawdown magnitudes of the Mediterranean Sea in the Levant basin during the Lago Mare stage of the Messinian Salinity Crisis[END_REF]Andreetto et al., 2021). These fluctuations are ascribed to switch in the Mediterranean freshwater budget driven by the African summer monsoon and Atlantic winter storms (e.g. [START_REF] Marzocchi | Orbital control on late Miocene climate and the North African monsoon: insight from an ensemble of sub-precessional simulations[END_REF][START_REF] Marzocchi | Precessional drivers of late Miocene Mediterranean sedimentary sequences: African summer monsoon and Atlantic winter storm tracks[END_REF]Simon et al., 2017). Since higher freshwater discharge rates occur at precession minima times and their Stage 3 sedimentary expression is considered to be the mudstone intervals (Fig. 3a; [START_REF] Manzi | A new facies model for the Upper Gypsum of Sicily (Italy): chronological and paleoenvironmental constraits for the Messinian salinity crisis in the Mediterranean[END_REF], mudstone interbeds (both onshore and offshore; e.g. Figs. 5h-j) represent the highstand episodes (e.g. [START_REF] Manzi | A new facies model for the Upper Gypsum of Sicily (Italy): chronological and paleoenvironmental constraits for the Messinian salinity crisis in the Mediterranean[END_REF]Roveri et al., 2008a;[START_REF] Omodeo-Salé | Tectonic and climatic control on the Late Messinian sedimentary evolution of the Nijar Basin (Betic Cordillera, Southern Spain)[END_REF]Fig. 3), while continental facies onshore (e.g. conglomerates in the Apennines; Fig. 5g) and offshore (clastic beds in the Levant Basin) and gypsum beds (Algero-Balearic, Liguro-Provencal, CMD, Tyrrhenian, Caltanissetta, Ionian, Sirte and Polemi-Pissouri basins; Figs. 5h-j) represent the lowstand (e.g. Roveri et al., 2008a;[START_REF] Manzi | A new facies model for the Upper Gypsum of Sicily (Italy): chronological and paleoenvironmental constraits for the Messinian salinity crisis in the Mediterranean[END_REF]Meilijson et al., 2019;Fig. 3). If Atlantic was the major source of sulfate for Stage 3 gypsum (e.g. García-Veigas et al., 2018) and an intervening, relatively shallow (Sicily) sill was establishing Western and Eastern Mediterranean division during the MSC (e.g. [START_REF] Garcia-Castellanos | Catastrophic flood of the Mediterranean after the Messinian salinity crisis[END_REF], 2020;[START_REF] Micallef | Evidence of the Zanclean megaflood in the eastern Mediterranean Basin[END_REF], the presence of Stage 3 gypsum to the east of the Sicily sill (Fig. 2b) implies that the Mediterranean base level never dropped below the (maximum estimated) paleodepth of the sill (i. e. ~430 m; [START_REF] Garcia-Castellanos | Catastrophic flood of the Mediterranean after the Messinian salinity crisis[END_REF] during Stage 3 and Western and Eastern Mediterranean remained connected also during the arid (lowstand) phases of the precession cycles. A Mediterranean step-wise refilled and at times filled with water up to the marginal belt agrees with: (1) Paratethyan biota being present only in intermediate-deeper settings during substage 3.1, but more widespread including marginal settings during substage 3.2; (2) the W-E homogeneity of Paratethyan ostracod assemblages around the Mediterranean marginal belt [START_REF] Gliozzi | Paratethyan ostracod immigrants in Italy during the Late Miocene[END_REF]Stoica et al., 2016;[START_REF] Sciuto | New reports of Paratethyan ostracods affinity from the Mediterranean Basin (Sicily, Italy)[END_REF][START_REF] Sciuto | Full restoration of marine conditions after the late Messinian Mediterranean Lago-Mare phase in Licodia Eubea and Villafranca Tirrena areas (east Sicily)[END_REF]; Fig. 9a); (3) the presence, in marginal basins, of Paratethyan fish [START_REF] Bannikov | Neogene Paratethyan croakers (Teleostei, Sciaenidae)[END_REF][START_REF] Schwarzhans | Late Miocene-Early Pliocene evolution of Mediterranean gobies and their environmental and biogeographic significance[END_REF], dinocysts (e.g. Pellen et al., 2017;Fig. 9b) and biomarkers (Vasiliev et al., 2017;Fig. 10c); (4) the occurrence of a monospecific assemblage of abundant Sphenolithus spp. just below the M/P boundary at ODP Sites 978, 975 and 967 [START_REF] Castradori | Calcareous nannofossils in the basal Zanclean of the Eastern Mediterranean Sea: remarks on paleoceanography and sapropel formation[END_REF]; [START_REF] Cita | Inventory of biostratigraphical findings and problems[END_REF] the requirement of water from the Mediterranean to explain the Sr isotope ratios of ostracods that inhabited marginal subaqueous environments (Andreetto et al., 2021); (6) the Atlantic-like sulfate values (although variably diluted and affected by microbial processes; Fig. 10b) of the UU/UG gypsum beds (García-Veigas et al., 2018); (7) the presence of long chain alkenones in the Sicilian UG beds similar to those observed in present-day marine settings (Fig. 10c; Vasiliev et al., 2017).

Major problems also exist with the half-full stratified scenario: (1) it does not provide a proper mechanism for gypsum precipitation at several hundreds, or thousands, meters water depth; (2) it fails to explain how unquestionable shallow-water (< 50 m) benthic organisms such as Ammonia tepida and Cyprideis sp. could survive at hundreds of meters of depth and beyond; (3) it does not provide an explanation for the high abundance of coarse-grained detritus at intermediate and deepwater locations, especially when compared to deep-water Pliocene samples, as well as for the broad absence of MSC deposits in the shelf domain; (4) a persistent Atlantic inflow without outflow seems to be a configuration that cannot be maintained stable for ~200 kyr. Indeed, models coupling the inflow of marine waters with the erosion of the gateway channel concluded that, if the Mediterranean level was lowered by at least several hundred meters below present sea level, any small overtopping of water from the Atlantic would inevitably trigger a fast refill of the basin that, if responsible for the erosion trough the Alborán Basin, should have involved an unprecedented water discharge and be completed within a few years or less (Garcia-Castellanos et al., 2020 and references therein). The scenario arising from Meilijson et al., 2019, Figs. 3b, 4a) is also problematic for a high base-level Mediterranean. In order to simultaneously reach precipitation of gypsum and halite in different basins sharing the same water, the water has to be of high salinity and stratified. Simon and Meijer (2017) demonstrated that this can be achieved with slow overturning circulation, but it is currently unclear how realistic this process is.

Demise of the MSC (5.33 Ma): the Zanclean marine replenishment

The conspicuous and abrupt transition to normal marine sediments in the Mediterranean is globally and historically important because it is the origin of the stratigraphic position of the Miocene/Pliocene boundary [START_REF] Van Couvering | The base of the Zanclean Stage and of the Pliocene Series[END_REF]. From an ocean perspective, it is not an ideal stratigraphic location being difficult to locate from biozone data even in the adjacent Atlantic [START_REF] Hodell | Correlation of late Miocene to early Pliocene sequences between the Mediterranean and North Atlantic[END_REF][START_REF] Andreetto | Revised astrochronology for the Ain el Beida section (Atlantic Morocco): no glacioeustatic control for the onset of the Messinian Salinity Crisis[END_REF][START_REF] Van Den Berg | Astronomical tuning for the upper Messinian Spanish Atlantic margin: disentangling basin evolution, climate cyclicity and MOW[END_REF]. However, from a Mediterranean perspective it provides a clear and unambiguous end to the MSC and the restoration of normal marine conditions. All evidence show that the onset of the Zanclean marine replenishment followed a period of relative lowstand that exposed all the Mediterranean margins (see subsection 3.12; Figs. 6a-b, f) and kept intermediate and deep basins underwater (see subsection 4.8). Yet again, the key dispute concerns the exact depth of the Mediterranean base level preceding the Miocene/Pliocene transition.

Building on the isolated Mediterranean scenario, base level immediately before the early Zanclean was more than thousand kilometers below eustatic sea level (Fig. 10e; e.g. Hsü et al., 1973a;[START_REF] Blanc | The opening of the Plio-Quaternary Gibraltar Strait: assessing the size of a cataclysm[END_REF][START_REF] Loget | Mesoscale fluvial erosion parameters deduced from modelling the Mediterranean sea-level drop during the Messinian (late Miocene)[END_REF][START_REF] Garcia-Castellanos | Catastrophic flood of the Mediterranean after the Messinian salinity crisis[END_REF][START_REF] Pérez-Asensio | Impact of restriction of the Atlantic-Mediterranean gateway on the Mediterranean Outflow Water and eastern Atlantic circulation during the Messinian[END_REF][START_REF] García-Alix | Updating the Europe-Africa small mammal exchange during the late Messinian[END_REF][START_REF] Amadori | Restored topography of the Po Plain-Northern Adriatic region during the Messinian base level drop-Implications for the physiography and compartmentalization of the paleo-Mediterranean basin[END_REF][START_REF] Micallef | Evidence of the Zanclean megaflood in the eastern Mediterranean Basin[END_REF][START_REF] Micallef | Geomorphic evolution of the Malta Escarpment and implications for the Messinian evaporative drawdown in the eastern Mediterranean Sea[END_REF][START_REF] Camerlenghi | Seismic markers of the Messinian salinity crisis in the deep Ionian Basin[END_REF][START_REF] Kartveit | Evidence of sea level drawdown at the end of the Messinian salinity crisis and seismic investigation of the Nahr Menashe unit in the northern Levant Basin, offshore Lebanon[END_REF][START_REF] Madof | Discovery of vast fluvial deposits provides evidence for drawdown during the late Miocene Messinian salinity crisis[END_REF][START_REF] Ben Moshe | Estimating drawdown magnitudes of the Mediterranean Sea in the Levant basin during the Lago Mare stage of the Messinian Salinity Crisis[END_REF][START_REF] Caruso | The late Messinian "Lago-Mare" event and the Zanclean Reflooding in the Mediterranean Sea: new insights from the Cuevas del Almanzora section (Vera Basin, South-Eastern Spain)[END_REF][START_REF] Cazzini | New seismic evidence of the Messinian paleomorphology beneath Lake Maggiore area (Italy)[END_REF][START_REF] Mas | The messinian salinity crisis in Mallorca: New insights for a western mediterranean stratigraphic scenario[END_REF][START_REF] Spatola | A single-stage megaflood at the termination of the Messinian salinity crisis: geophysical and modelling evidence from the eastern Mediterranean Basin[END_REF]. Hydrodynamic erosional models allowed a reinterpretation of the erosional features across the strait of Gibraltar [START_REF] Campillo | Stratigraphic and Tectonic evolution of the Western Alboran Sea late Miocene to recent[END_REF][START_REF] Blanc | The opening of the Plio-Quaternary Gibraltar Strait: assessing the size of a cataclysm[END_REF] suggesting that a sudden breach of the Mediterranean-Atlantic divide at Gibraltar resulted in a vast cascade of Atlantic water that refilled the entire Mediterranean in less than 2 years (i.e. rates of ten meters per day) spilling first into the Western Mediterranean (see the extensive review in Garcia-Castellanos et al., 2020) and then, after reaching the level of the Sicily sill, pouring into the Eastern Mediterranean [START_REF] Micallef | Evidence of the Zanclean megaflood in the eastern Mediterranean Basin[END_REF][START_REF] Micallef | Geomorphic evolution of the Malta Escarpment and implications for the Messinian evaporative drawdown in the eastern Mediterranean Sea[END_REF][START_REF] Ben Moshe | Estimating drawdown magnitudes of the Mediterranean Sea in the Levant basin during the Lago Mare stage of the Messinian Salinity Crisis[END_REF][START_REF] Spatola | A single-stage megaflood at the termination of the Messinian salinity crisis: geophysical and modelling evidence from the eastern Mediterranean Basin[END_REF]. This concept of catastrophic refilling has led to terms such as "Zanclean flood" or "deluge". Evidence supporting the catastrophic flood mechanism mostly comes from the seismic reflection dataset and includes: 1) the presence of >250 m deep and 390-km-long incisions on both sides of the Gibraltar Strait [START_REF] Garcia-Castellanos | The Zanclean megaflood of the Mediterranean-Searching for independent evidence[END_REF]; 2) the detection of (allegedly) Pliocene-aged chaotic sedimentary bodies stretching for kilometers in the Alborán Basin (Garcia-Castellanos et al., 2020 and references therein) and Ionian Basin at the foot of the Malta Escarpment [START_REF] Micallef | Evidence of the Zanclean megaflood in the eastern Mediterranean Basin[END_REF][START_REF] Micallef | Geomorphic evolution of the Malta Escarpment and implications for the Messinian evaporative drawdown in the eastern Mediterranean Sea[END_REF][START_REF] Spatola | A single-stage megaflood at the termination of the Messinian salinity crisis: geophysical and modelling evidence from the eastern Mediterranean Basin[END_REF]Fig. 8c). A further argument is the bathymetric jump of several hundred meters between the late Messinian and the early F. Andreetto et al. Pliocene sediments (e.g. [START_REF] Caruso | The late Messinian "Lago-Mare" event and the Zanclean Reflooding in the Mediterranean Sea: new insights from the Cuevas del Almanzora section (Vera Basin, South-Eastern Spain)[END_REF]Fig. 6d).

Instantaneous sea level rise is not the only possible refilling model. [START_REF] Bache | A twostep process for the reflooding of the Mediterranean after the Messinian Salinity Crisis[END_REF] suggested the reflooding occurred in two steps at ~5.60 Ma, accompanied by a moderate (≤ 500 m) rise, followed by a rapid rise of 600-900 m at around 5.46 Ma tracking the deposition of the deep basin evaporites and resulting from the collapse of the Gibraltar divide. There is also the reconnection model that follows from a Stage 3 Mediterranean that is already relatively full and with the base level possibly oscillating of 400 ± 100 m with precessional frequency (Fig. 10h; [START_REF] Fortuin | The Messinian of the Nijar basin (SE Spain): sedimentation, depositional environments and paleogeographic evolution[END_REF][START_REF] Ben Moshe | Estimating drawdown magnitudes of the Mediterranean Sea in the Levant basin during the Lago Mare stage of the Messinian Salinity Crisis[END_REF]Andreetto et al., 2021). In this case, only a sea level rise of a few hundred meters is required to restore the Mediterranean to the Atlantic level (Fig. 10h), which was hypothesized to have occurred in the last precessional cycle of the Messinian (Marzocchi et al., 2016;Fig. 3a).

In detail, the re-establishment of a fully marine faunal diversity and oceanic geochemistry (e.g. 87 Sr/ 86 Sr ratios and δ 18 O) occurred more gradually over one or more precessional cycles in the earliest Zanclean (e.g. Iaccarino et al., 1999;[START_REF] Pierre | Sedimentological and stable isotope changes at the Messinian-Pliocene boundary in the eastern Mediterranean[END_REF][START_REF] Pierre | Reconstruction of the paleoenvironmental changes around the Miocene-Pliocene boundary along a West-East transect across the Mediterranean[END_REF][START_REF] Cipollari | Easternmost Mediterranean evidence of the Zanclean flooding event and subsequent surface uplift: Adana Basin, southern Turkey[END_REF]Roveri et al., 2019a;Bulian et al., 2021). This suggests that stressed ecological conditions at first only suitable for opportunistic organisms to survive (e.g. Bulian et al., 2021) developed (or persisted) in the Mediterranean as marine replenishment occurred (e.g. [START_REF] Rouchy | Late Messinian to Early Pliocene paleoenvironmental changes in the Melilla Basin (NE Morocco) and their relation to Mediterranean evolution[END_REF]. One possible mechanism for achieving this may be the physicochemical turnover in the water column triggered by the re-established two-way exchange with the Atlantic which, for reasons that are largely unknown, took time (at least oneprecession cycle; [START_REF] Pierre | Reconstruction of the paleoenvironmental changes around the Miocene-Pliocene boundary along a West-East transect across the Mediterranean[END_REF] to displace surficial Paratethyan water and restore normal marine conditions (Marzocchi et al., 2016).

Methods and proxies to better reconstruct base level and connectivity changes

Chronological uncertainty and spatial variability limit the use of both sedimentological and paleontological information to achieve a comprehensive and coherent basin-wide interpretation of the conditions and drivers of Stage 3 environments and water levels. Alternative methods are therefore required to clarify connectivity relationships and constrain base-level conditions. This section explores the principles and potential of geochemical, backstripping and numerical modelling techniques that could be used to further test existing hypotheses and enhance understanding of the complex environmental conditions experienced by the Mediterranean during the latest Messinian.

Geochemical proxies

Radiogenic strontium isotopes. Radiogenic strontium isotope ratio ( 87 Sr/ 86 Sr) is a widely applied geochemical tool in provenance studies, including the reconstruction of the hydrological circuit and connectivity of basins with little or null oceanic entries. Its potential to detect the provenance of the hydrological fluxes derives from the unique 87 Sr/ 86 Sr ratio that typifies each water source and from the negligible effects of isotopic fractionation during the liquid-solid transition (see [START_REF] Hajj | Radiogenic and "stable" strontium isotopes in provenance studies: a review and first results on archaeological wood from shipwrecks[END_REF].

Mineral phases precipitating in endorheic lakes uptake Sr with 87 Sr/ 86 Sr ratio that reflects the mixing of all feeding surficial and underground streams and whose 87 Sr/ 86 Sr fingerprint hinges on the composition and age of watershed bedrock (see [START_REF] Peucker-Ehrenbrink | A continental perspective of the seawater 87Sr/86Sr record: A review[END_REF]Andreetto et al., 2021 and references therein). When river water mixes with seawater such as in the oceans, semi-enclosed basins or estuaries, mineral phases uptake Sr with oceanic 87 Sr/ 86 Sr ratios because the high oceanic Sr concentration (~7.8 mg/l today; [START_REF] Veizer | Strontium isotopes in seawater through time[END_REF] masks the impact of the ~100 times lower concentrated continental Sr-sources (~0.0780 mg/l; [START_REF] Palmer | Controls over the strontium isotope composition of river water[END_REF]. This is valid as long as a certain ratio of continental-marine water mixing is fulfilled, beyond which 87 Sr/ 86 Sr ratios deviate towards the 87 Sr/ 86 Sr ratios of the non-marine source(s) [START_REF] Ingram | Strontium isotopic composition of estuarine sediments as paleosalinity-paleoclimate indicator[END_REF]. For the Mediterranean to attain non-marine 87 Sr/ 86 Sr ratios (like during the MSC), Topper et al. (2014) calculated a mixing of at least 1:4 (Atlantic: freshwater) to be required.

If Mediterranean subbasins hosted endorheic lakes (Figs. 10c, e), the 87 Sr/ 86 Sr isotope ratios measured on ostracod valves or gypsum crystals of each lake are expected to generate a scattered distribution by virtue of the different geology in the hinterland of each basin. By contrast, some degree of connection between different basins and the Mediterranean water mass (Figs. 10d,f) is expected to result in more homogeneous 87 Sr/ 86 Sr ratios because, although isotopically-different, local rivers mix with a water mass that has the same 87 Sr/ 86 Sr value and (much higher) Sr concentration for each basin (Andreetto et al., 2021). In this scenario, differences in the 87 Sr/ 86 Sr ratios between basins are likely the reflection of the different 87 Sr/ 86 Sr ratio of the local input in each basin (Andreetto et al., 2021). The application of numerical models assists to identify and quantify the different water sources feeding the basin(s) in question and (e.g. [START_REF] Placzek | Isotopic tracers of paleohydrologic change in large lakes of the Bolivian Altiplano[END_REF]Topper et al., 2011Topper et al., , 2014;;[START_REF] Doebbert | Controls on Sr isotopic evolution in lacustrine systems: Eocene green river formation[END_REF]Rossi et al., 2015b;[START_REF] Modestou | Precessional variability of 87 Sr/ 86 Sr in the late miocene sorbas basin: An interdisciplinary study of drivers of interbasin exchange[END_REF]Grothe et al., 2020;Andreetto et al., 2021).

Sulfate isotopes. When sulfate-bearing minerals precipitate in a basin they uptake dissolved S and O with δ 34 S SO4 and δ 18 O SO4 isotopic composition that, once corrected for the fractionation effects during liquid-solid transition (see subsection 6.2), can be assimilated to that of the mother brine. The higher concentrated source of sulfate is seawater (with present-day δ 34 S SO4 =21.15±0.15‰ and δ 18 O SO4 =8.67±0.21‰, [START_REF] Johnston | Placing an upper limit on cryptic marine sulphur cycling[END_REF] with Messinian values of ~22±0.2‰ for the δ 34 S SO4 and ~9±2‰ for the δ 18 O SO4 ; Turchyn and Schrag, 2004;Markovic et al., 2016;Masterson et al., 2016). Significantly higher inputs from the ~1000 times less concentrated riverine freshwater (both surficial and underground) with respect to the ocean water (more than 1:5 according to Lu et al., 2001) can modify the marine δ 34 S SO4 and δ 18 O SO4 isotopic composition of the mother brine (Utrilla et al., 1992;Lu et al., 2001) and have it deviated from that of the ocean (Lu et al., 2001). This deviation is normally towards lower values, because river-derived dissolved sulfate is generally depleted in heavy isotopes 34 S and 18 O compared to oceanic sulfate because these isotopes mainly come from the oxidation of 34 S-depleted pyrite (FeS 2 ) on the continents and to a lesser extent from the dissolution of older sulfate-bearing minerals [START_REF] Claypool | The age curves of sulfur and oxygen isotopes in marine sulfate and their mutual interpretation[END_REF]Turchyn and Schrag, 2004;Burke et al., 2018). However, when marine sulfate is preferentially leached in the catchment, 34 S of the freshwater-dissolved sulfate and [SO 4 2-] likely increase, therefore reducing the continental-marine mixing ratio necessary to deviate the resulting sulfate isotopic signature away from marine values. Unfortunately, the sulfate isotopic composition is not provided for a number of major Mediterranean rivers (Burke et al., 2018) nor for the Eastern Paratethys and it is hardly assessed with the catchment-forming lithologies [START_REF] Liu | Sulfate sources constrained by sulfur and oxygen isotopic compositions in the upper reaches of the Xijiang River, China[END_REF]Burke et al., 2018), making sulfate isotopes still an unsuitable tracer of non-marine water provenance in Mediterranean subbasins.

Deviation of δ 34 S SO4 and δ 18 O SO4 from the marine average can also be the result of isotopic fractionation during microbial sulfate reduction (MSR; Fritz et al., 1989;[START_REF] Berner | Atmospheric oxygen over Phanerozoic time[END_REF]. MSR produces 34 S-depleted hydrogen sulfide (~0 to 70‰ lighter than initial sulfate; Brunner and Bernasconi, 2005;Sim et al., 2011;Leavitt et al., 2013) and induces the enrichment in 34 S and 18 O of the residual sulfate pool [START_REF] Kaplan | Microbiological fractionation of sulphur isotopes[END_REF]Thode and Monster, 1965;[START_REF] Turchyn | Oxygen isotopic composition of sulfate in deep sea pore fluid: evidence for rapid sulfur cycling[END_REF]Wortmann et al., 2007). Therefore, if isotopically light H 2 S produced by MSR leaves the system as a sulfide mineral (most likely pyrite), the resulting dissolved sulfate would have δ 34 S SO4 and δ 18 O SO4 isotopic signatures higher than the oceanic one (Brunner et al., 2005). However, if the MSRproduced H 2 S is re-oxidized back to sulfate through abiotic or microbial sulfide oxidation, isotopically light sulfate will be brought back to the 34 S-enriched sulfate pool, producing little or no enrichment in 34 S observed in the resulting sulfate (Gomes and Johnston, 2017 and references therein;Pellerin et al., 2019). Slight deviations from marine δ 18 O SO4 and δ 34 S SO4 values of sulfate reflect both biological sulfur F. Andreetto et al. cycling and/or freshwater riverine inputs (e.g. Utrilla et al., 1992;Lu et al., 2001;[START_REF] Turchyn | Stable isotope analysis of the Cretaceous sulfur cycle[END_REF] (Fig. 10b). Untangling the relative importance of these processes is key to understanding the Mediterranean sulfur isotope record and gleaning paleoenvironmental insights into Stage 3.

Hydrogen isotopes. Organic geochemistry biomarker-based tools can be used as independent proxies for reconstructing sea surface temperatures, relative changes in the basin hydrology and, indirectly, salinity. Basin water properties are reflected in a variety of life forms. Different types of organisms produce specific organic compounds that serve as molecular biomarkers. These large biomolecules record the changes in the hydrogen isotopic composition of the water used by different groups of biomarker producers (i.e. different organisms). The principle behind the method is to measure δD on biomarkers produced in Mediterranean Sea waters (e.g. alkenones, produced by a few species of haptophyte coccolithophores algae) during the MSC and compare the results with the δD signals retrieved from biomarkers produced in the open ocean ideally at the same time intervals. The influence of precipitation on the changes in hydrological budget can be monitored by measuring the δD of long chain n-alkanes [START_REF] Sachse | δD values of individual n-alkanes from terrestrial plants along a climatic gradient -implications for the sedimentary biomarker record[END_REF], biomarkers predominantly produced by higher terrestrial plants that rely on precipitation for plant growth, therefore reflecting the changes in the δD of the precipitation. The extreme base level drop(s) suggested for the Mediterranean during the MSC would, in principle, indicate a negative precipitation (P) + runoff (R) -evaporation (E) ratio. Such a negative water budget (E>P+R) results in waters increasingly enriched in δD whereas, a positive water balance (E<P+R) results instead in a negative shift of δD values. The analysis of compound specific δD of alkenones, long and short chain n-alkanes can be used to constrain E/(P+E) relationships.

Backstripping analyses

Backstripping uses paleobathymetry, sea level and sediment thickness to quantify the tectonic and isostatic components of subsidence. If tectonic subsidence or uplift history are known relative to the current position and depth of paleoshoreline markers, an inverse approach allows base level to be estimated. A number of approaches have been applied to the MSC, using erosional surfaces (e.g. [START_REF] Amadori | Restored topography of the Po Plain-Northern Adriatic region during the Messinian base level drop-Implications for the physiography and compartmentalization of the paleo-Mediterranean basin[END_REF], terraces [START_REF] Micallef | Evidence of the Zanclean megaflood in the eastern Mediterranean Basin[END_REF] or fluvial network characteristics [START_REF] Urgeles | New constraints on the Messinian sealevel drawdown from 3D seismic data of the Ebro margin, western Mediterranean[END_REF] as paleoshoreline indicators. The relief on erosional features has also been used to estimate minimum base-level variation [START_REF] Frey-Martinez | 3D seismic interpretation of the Messinian Unconformity in the Valencia Basin, Spain[END_REF].

Apart from the quantitative constraints on base level that backstripping provides, consideration of the regional implications of isostatic subsidence and the gravitational impact of redistributing water masses (such as in the cascading model of Roveri et al., 2014c;Fig. 10b) and evaporite precipitation is important in gateway regions like Gibraltar, which due to their shallow and restricted nature are exceptionally sensitive to vertical motions. Here, both flexural effects and gravitational effects on local sea level on the Atlantic side of the strait has the potential to influence Mediterranean-Atlantic connectivity driving paleoenvironmental changes in the basin itself [START_REF] Coulson | The role of isostatic adjustment and gravitational effects on the dynamics of the Messinian salinity crisis[END_REF].

Modelling

Numerical models can be used complementary to lab-and fieldbased studies, or to find answers to open questions by testing the physical plausibility of hypotheses and their compatibility with the available sedimentological/stratigraphic/paleontological/geochemical data, which have to constrain model results and not adjust to it. For example, whether gypsum beds in marginal/intermediate basins can precipitate at the same time as the halite in deep basins is an intriguing question that circulates in the MSC literature (e.g. [START_REF] Van Couvering | The terminal Miocene event[END_REF], but whether this is physically and geochemically possible is yet to be answered. In a first model analysis, Simon and Meijer (2017) found that the required stratification can indeed be achieved for specific conditions including a slow overturning circulation. A different approach is needed to determine whether such slow circulation is to be expected or if other scenarios should be considered. A thermo-haline stratification that enables coeval precipitation of two evaporites for a considerable time span might also influence the degree of heterogeneity of other parameters, such as strontium concentration. Previous models showed the influence of the Atlantic Ocean and major rivers (Topper et al., 2014) and of evaporation (Flecker et al., 2002) on the Sr value of a basin with restricted oceanic inflow and assuming a homogeneous distribution of strontium throughout the water column (Flecker et al., 2002;Topper et al., 2011Topper et al., , 2014;;[START_REF] Modestou | Precessional variability of 87 Sr/ 86 Sr in the late miocene sorbas basin: An interdisciplinary study of drivers of interbasin exchange[END_REF], but it is uncertain if this holds true in conditions of water stratification. New insights into this behavior would have consequences for the way the strontium dataset (Fig. 10a) must be interpreted. Another loose end where the model approach can provide insight relates to the question whether a high water level could have been reached without an inflow from the Atlantic. Climate simulations conducted by [START_REF] Gladstone | The Mediterranean hydrologic budget from a Late Miocene global climate simulation[END_REF], Simon et al. (2017) and [START_REF] Marzocchi | Precessional drivers of late Miocene Mediterranean sedimentary sequences: African summer monsoon and Atlantic winter storm tracks[END_REF] indicate that this is not possible with today's bathymetry. A quantitative analysis exploring the Mediterranean water level reached in different situations (i.e. with or without an Atlantic or Paratethys in and outflow) and with information on the Mediterranean hypsometry that may be provided by isostatic restoration of the seafloor topography (flexural backstripping) could help understanding how the Messinian Salinity Crisis ended.

Certainties, open problems and future directions

Our understanding of the nature of MSC Stage 3 has evolved considerably over the last fifty years. However, there are still such disparate models for the paleoenvironmental conditions and basin connectivity that led to Stage 3 deposition and that express the challenges that the study of this interval presents: it is a relatively short interval and its sedimentary expression varies spatially. It is no surprise that the main point of contention lies in reconciling the observations from seismic profiles and well data, which are largely interpreted as favoring the desiccated scenario, with the sedimentological, paleontological and geochemical data from the marginal basins record, largely discontinuous and unaddressed from seismic-based and computationalbased studies and mainly supporting the full-Mediterranean hypothesis.

To a large extent this mismatch is the result of the lack of intersection of the two datasets. Some Stage 3 onshore localities are meticulously studied from the stratigraphic, sedimentological, paleontological and geochemical point of view, showing remarkably consistent and homogenous trends and patterns (as previously highlighted by Roveri et al., 2008a). However, changes at precessional and subprecessional scale are difficult to trace from one exposure to another and are well below the level of seismic resolution, making onshore-offshore correlation at this scale impossible. Even correlation between onshore sections is problematic since most of the stratigraphic sections are incomplete, with erosion surfaces at the bottom and/or top (i.e. the Miocene/Pliocene boundary), and this lack of stratigraphic continuity frustrates attempts to constrain ages by cyclostratigraphy. A future focus on strengthening the available chronostratigraphic framework (Fig. 3) and making it inclusive of the fragmented outcrops is required to better understand the paleoenvironmental and paleohydrological changes suffered by the Mediterranean marginal belt through time. The successful drilling of the three IODP proposals currently in the scheduling pool (see [START_REF] Camerlenghi | Uncovering the Mediterranean Salt Giant (MEDSALT)-Scientific Networking as Incubator of Cross-disciplinary Research in Earth Sciences[END_REF], all of which propose to recover Stage 3 sediments, will also provide transformative information enabling far better offshore-onshore correlation and interpretation of currently enigmatic seismic data. In the meantime, re-evaluation of existing DSDP and ODP material, particularly through the application of more novel geochemical techniques and, where possible, access to material collected during industrial drilling would be helpful for understanding the deep Mediterranean Basin during this interval.

Extensive paleontological studies have established that Stage

APPENDIX D

contains in situ biota assemblages of Paratethyan provenance implying brackish water conditions. More problematic is the differentiation of in situ and reworked marine microfauna and flora and the paleoecological significance of dwarfism in marine calcareous microfossils/algae. These have important repercussions for the Mediterranean connectivity and base-level reconstruction.

The geochemical dataset for Stage 3, particularly strontium isotopes and hydrogen isotopes on biomarkers, is both demonstrably valuable in providing key constraints on connectivity and environmental conditions, and frustratingly inadequate in terms of data distribution. It has great potential as a constraint on quantitative sensitivity analysis of the different hydrochemistry scenarios that follow from the endmember Stage 3 hypotheses, but substantially more data is required.

An approach which combines geological, geochemical, geophysical and paleontological data with numerical modelling (e.g. climate simulations, backstripping analyses and paleoceanographic models) will provide a more accurate reconstruction of Mediterranean paleogeography and the processes that occurred during the final phase of the Messinian Salinity Crisis.

The formation and dissolution of salt giants impacts ocean chemistry on thousand-million year timescales. Gypsum precipitation and weathering changes the oceanic calcium concentration with implications for the carbon cycle and global temperatures. However, the connectivity of salt giants with the global ocean is necessarily restricted, making the timing of Ca 2+ extraction and return more uncertain. Here we reconstruct the final phase of gypsum precipitation of the Late Miocene Mediterranean Salt Giant using micropaleontology, sedimentology and 87 Sr/ 86 Sr analyses on the most complete record preserved at Eraclea Minoa on Sicily and explore its implications for global climate. Precessional gypsum-marl couplets (Upper Gypsum) characterize the last 200 kyrs (Stage 3) of the Messinian Salinity Crisis (MSC; 5.97-5.332 Ma) in both intermediate (500-1000 m) and deep (>1000 m) Mediterranean basins. The interbedded selenitic gypsum layers contain well-preserved calcareous nannofossil assemblages dominated by Reticulofenestra minuta, a marine species which tolerates stressful conditions. Marine water is also required to explain the gypsum 87 Sr/ 86 Sr data, which describe a small range of ratios (0.708704-0.708813) lower than coeval ocean water. Mass-balance calculations indicate that during gypsum precipitation, the Atlantic made up ≤20% of a Mediterranean ("Lago-Mare") water mass dominated by low salinity discharge from large river systems and Eastern Paratethys. This suggests episodic extraction of calcium and sulfate ions from the ocean throughout MSC Stage 3. The marls commonly contain shallow (30-100 m) brackish-water ostracods of Paratethyan (Black Sea) origin. Marls with Paratethyan ostracods are also found in both marginal (<500 m) and deep Mediterranean settings. This indicates that marl-deposition was not synchronous across the basin, but that it occurred in intermediate and deep basins during base-level lowstands at insolation minima and on the shallow Mediterranean margins during insolation maxima-driven highstands. These high-amplitude base-level fluctuations exposed the evaporites to weathering, but ponded the products in the Mediterranean basin until reconnection occurred at the beginning of the Pliocene.

Introduction

Recent modeling of evaporite weathering and deposition suggests that, alongside the well-established silicate weathering and carbonate sedimentation feedback mechanisms, the formation and dissolution of giant calcium sulfate deposits may be an episodic driver of thousand-million year timescale changes in the carbon cycle (Shields and Mills, 2020). Salt giants have formed only episodically in Earth history, typically as a consequence of the tectonic restriction (opening or closing) of the connection between the global ocean and a mid-latitude marginal marine basin (Warren, 2016). Model experiments simulate the atmospheric CO 2 and temperature consequences of evaporite deposition and weathering events assuming a constant oceanic supply of Ca 2+ to the marginal basins, or a constant riverine Ca 2+ and SO 2- 4 flux from continental weathering, both lasting an estimated 1 million years (Shields and Mills, 2020). The heterogenous geological record of these salt giants, however, suggests that the ocean-marginal basin connectivity must have varied considerably (Warren, 2016), impacting the duration and intensity of ion fluxes in relation to evaporite precipitation and weathering during the formation of a salt giant. Reconstructing, the nature and timing of that variable connectivity, its relationship with gypsum precipitation and baselevel changes is a critical step in quantifying the importance of this mechanism in driving carbon cycle change.

The most recent of the Earth's salt giants accumulated in the Mediterranean during the Messinian Salinity Crisis (MSC; . This formed as a consequence of Late Miocene tectonic uplift of the Gibraltar region which restricted marine exchange between the Mediterranean Basin and the Atlantic Ocean (e.g. Krijgsman et al., 2018). The Mediterranean's environmental response was dramatic, resulting in the precipitation of >1 million km 3 of evaporites (e.g. Roveri et al., 2014a). Three evaporitic episodes with different connectivity configurations have been distinguished (Roveri et al., 2014a): Stage 1, dominated by gypsummarl couplets which are thought to have been deposited during restricted two-way Atlantic-Mediterranean exchange (e.g. Reghizzi et al., 2018); Stage 2 comprises an up to 2k m thick halite layer resulting from Atlantic inflow but no Mediterranean outflow (e.g. Krijgsman et al., 2018); and Stage 3 in which gypsum alternates with brackish-water, ostracod-bearing marls where the connectivity scenario is controversial (Andreetto et al., 2021a). It has proved difficult to explain both the outcrop data, which mostly suggest a relatively full Mediterranean, and evidence from deep-sea cores and seismics, which suggest low water conditions (see Andreetto et al., 2021a for summary). These different base level reconstructions have implications for Atlantic-Mediterranean connectivity at the end of the MSC and its impact on ocean chemistry.

Previous attempts to reconstruct Mediterranean hydrology, paleoenvironments and connectivity during Stage 3 have involved 87 Sr/ 86 Sr (Müller et al., 1990;Roveri et al., 2014b;García-Veigas et al., 2018) because water sources of different origin (i.e. ocean, lakes, rivers, groundwater) have distinct 87 Sr/ 86 Sr ratios (e.g. Reghizzi et al., 2018;Andreetto et al., 2021bAndreetto et al., , 2022;;[START_REF] Bista | Sr isotope-salinity modelling constraints on Quaternary Black Sea connectivity[END_REF]. A solely geochemical approach to Stage 3 has proved inconclusive however, because more than one combination of potential source waters can result in the same geochemical signal. In this multi-proxy study, we combine new and existing sedimentological and paleontological observations with 87 Sr/ 86 Sr analyses on selenite gypsum crystals and ostracod valves from the Eraclea Minoa section (Caltanissetta Basin, Sicily; Fig. 1).

Eraclea Minoa is the most complete sedimentary record of Stage 3 exposed on land and previous works have described evidence of both Atlantic and Eastern Paratethys influence (Grossi et al., 2015;Vasiliev et al., 2017;García-Veigas et al., 2018). The succession shows strong precessional cyclicity from which an astronomical age model has been derived (van der Laan et al., 2006;[START_REF] Manzi | A new facies model for the Upper Gypsum of Sicily (Italy): chronological and paleoenvironmental constraits for the Messinian salinity crisis in the Mediterranean[END_REF]. Unlike age-equivalent marginal exposures, which are dominated by continental clastics, the Caltanissetta Basin was, during Stage 3, at intermediate depths (500-1500 mbsl) and accumulated both gypsum and ostracod-bearing marls, allowing the Sr isotopic characterization of both lithological elements. This dataset allows us to reconstruct Stage 3 connectivity and consider its impact on Late Miocene ocean biogeochemistry and climate.

Materials and methods

The Eraclea Minoa section (37 • 23'36"N, 13 • 17'10.38"E) is ∼223 m-thick and comprises seven gypsum-marl alternations (Fig. 2). This Upper Gypsum succession is overlain sequentially by the Are-nazzolo Fm., a ∼6-7 m-thick fine to medium-grained sandstone, and the Pliocene Trubi marls (Fig. 2). Field work to log and sample the Eraclea Minoa section was undertaken in Spring 2001 and September 2020 and 2021.

Selenitic gypsum can preserve biomineralized structures (Pellegrino et al., 2021), so twenty selenite crystals were selected from gypsums II to VII for micropaleontological screening. Crystals were cleaned of surface dirt using a brush and MilliQ purified water, sliced with a cleaned box cutter into 1-2 mm-thick splits and qualitatively analyzed, prior to carbon-coating the slices. Samples were then analyzed at the Electron Microscopy Centre at Utrecht University using a JEOL Neoscope II JCM-6000 table-top SEM equipped with an energy-dispersive X-ray spectroscopy (EDS) and Electron Probe Microanalyser JEOL JXA 8530F.

We expanded the available ostracod and gypsum 87 Sr/ 86 Sr dataset at Eraclea Minoa (Grossi et al., 2015;García-Veigas et al., 2018) with 43 additional isotopic ratios on gypsums (39) and Cyprideis agrigentina valves from cycle 3 (4). Sample preparation and data acquisition followed standard procedures for gypsum (Reghizzi et al., 2018) and ostracods [START_REF] Bista | Sr isotope-salinity modelling constraints on Quaternary Black Sea connectivity[END_REF]Andreetto et al., 2022). The more vitreous selenites and balatino gypsum chips were mechanically isolated, washed with MilliQ-water and micromilled with a hand drill avoiding impurities trapped within crystals. 10 mg of gypsum powder was digested in 1M ammonium acetate, centrifuged with MilliQ-water, dried overnight at 120 • C on a hot plate, redissolved in 3M HNO 3 , centrifuged and refluxed on the hot plate. 6-7 adult valves of C. agrigentina were hand-picked under an optical microscope, repeatedly rinsed with MilliQ-water and methanol to remove potential contaminants (clay particles and organic material), dissolved in 300 µl of 0.1 M of HNO 3 and dried at 120 • C. Strontium of all samples was isolated using 100 µl of pre-cleaned Sr-Spec resin and loaded onto rhenium filaments with a TaCl 5 activator solution. 87 Sr/ 86 Sr ratios from gypsums V and VI were measured in SUERC (University of Glasgow) on a VG Sector 54-30 mass spectrometer in dynamic multi-collection mode. Instrumental mass fractionation was corrected using an exponential law and 86 Sr/ 88 Sr=0.1194. NIST SRM987 values over the period the measurements were carried out were: 87 Sr/ 86 Sr =0.710251T20 (2cS.D.), n=95. 87 Sr/ 86 Sr ratios from the ostracods and gypsums I-IV and VII were measured at the Bristol Isotope Group Facilities (University of Bristol) on a Thermo-Finnigan Triton thermal ionization mass spectrometer using a multi-dynamic 'triple jump' method. The average 87 Sr/ 86 Sr is based on 200 ratios. The mean 87 Sr/ 86 Sr values of nine NIST SRM 987 analyzed during this session was 0.710245±0.000006267 and that of four NIST SRM 987 which have undergone chemical purification (mean 0.710268±0.0000236925) are in good agreement with the accepted values of SRM 987 of 0.710250. andC) and a single gypsum-free facies association (A). On the basis of more detailed lithological and faunal analysis, we divide their facies association A into two sub-units A and D.

Results

Sedimentological description of a typical UG cycle

Our sub-unit A, arbitrarily chosen as the base of the lithological cycle, comprises cm-to dm-thick beds of greenish, homogeneous marls interbedded with white laminated diatomites (Figs. 3c-d). This sub-unit, typically 1-2 meters thick, occurs immediately beneath the gypsum-bearing sequence (Figs. 2, 3b-c). Sub-unit D, at the top of the cycle, is found directly above the gypsum and consists of gray marls intercalated with cm-to dm-thick siltstones and 2 litharenites (Figs. 2,3h). The gypsum layers at Eraclea Minoa typically start with cm-to dm-thick beds of laminar (or balatino) and clastic gypsum (gypsarenites) alternating with greenish homogeneous marls (sub-unit B; Figs. 3c,e) and show a sharp change to rhythmic alternations of dm-thick layers of massive selenite and mm-to cm-thick greenish marls (sub-unit C; Fig. 3f). The balatino gypsum comprises alternating mm-thick gypsum cumulate layers with clay-rich horizons only a few hundreds of microns thick (Fig. 3e). The gypsarenites are commonly massive (Fig. 3e) with occasional sedimentary structures such as normal grading and horizontal lamination [START_REF] Manzi | A new facies model for the Upper Gypsum of Sicily (Italy): chronological and paleoenvironmental constraits for the Messinian salinity crisis in the Mediterranean[END_REF]. The selenite beds comprise bottom-grown, swallow-tail gypsum crystals (or selenites) oriented with the re-entrant angle of the twins upward (Fig. 3f). Crystal size is relatively uniform in individual beds, but increases vertically through sub-unit C, from cm-scale crystals at the base to several dm in length at the top. The geometry of the gypsum beds also changes upwards, from tabular (Fig. 3f) to domal or cauliflower-shaped (Fig. 3b). The resulting irregular upper surface (Fig. 3b) led some authors to interpret an episode of subaerial erosion (e.g. Rouchy and Caruso, 2006;Raad et al., 2021). However, no karstic features are observed in the Eraclea Minoa gypsums (see also [START_REF] Manzi | A new facies model for the Upper Gypsum of Sicily (Italy): chronological and paleoenvironmental constraits for the Messinian salinity crisis in the Mediterranean[END_REF].

At Eraclea Minoa, cycles 5, 6 and 7 all show a consistent facies sequence from the diatomite-bearing sub-unit A, through the cumulate and clastic gypsum sub-unit B, selenite-dominated subunit C, to the ostracod-bearing marls of sub-unit D (Fig. 2). Cycles 3 and 4 comprise sub-units B-D without the diatomaceous subunit A. Cycles 1 and 2 are poorly exposed with only sub-unit B visible in cycle 1 and C and D in cycle 2 and 3 (Fig. 2).

Micropaleontology

Ostracods in sub-unit D marls

Sub-unit D marls contain 25 ostracod species (Grossi et al., 2015). One of these, C. agrigentina, is of Mediterranean origin (Grossi et al., 2015). All the others are endemic to Paratethys (Fig. 2; Stoica et al., 2016). The absence of Paratethyan ostracods in Mediterranean sediments prior to MSC Stage 3 and the occurrence here of well-preserved valves of different moult stages indicate that these ostracods are in situ (Grossi et al., 2015). In cycles 3, 6 and 7 there is a clear increase in ostracod diversity through sub-unit D from almost (cycle 7) or completely (cycle 3 and 6) monospecific assemblages of C. agrigentina (Biofacies 1) at the base to diverse communities of Paratethyan taxa (Biofacies 2) at the top (Fig. 2; Grossi et al., 2015). Planktic and benthic foraminifera have also been reported from sub-unit D's marls [START_REF] Caruso | The Upper Gypsum unit[END_REF]. However, only Ammonia tepida, which like C . agrigentina dwells in euryhaline and shallow-water, is considered to be in-situ (Grossi et al., 2015). One of the three samples from a marl layer interbedded with selenites (sub-unit C) also contained valves of C. agrigentina and Loxoconcha muelleri (Fig. 2; Grossi et al., 2015). Analysis of a small number of marl samples from sub-unit A suggests that these sediments contain no ostracods (Fig. 2).

Calcareous nannofossils in selenitic gypsum

The Eraclea Minoa selenites display alternating turbid/dark and transparent laminae (Fig. 4a). SEM inspection of samples from layers II-VII reveals that the dark color is associated with abundant, closely packed calcareous nannofossils associated with a small detrital component, mostly clay and rare silt-and clay-sized euhedral halite crystals (Fig. 4b), which might result from drying out of fluid inclusions. Since the analyzed samples are slices of selenite crystals, we can confirm that these are microfossils embedded within the gypsum (Figs. 4b-d). Well preserved calcareous nannofossils comprising pristine platelets and intact coccospheres (Fig. 4e) are observed both in cavities (Figs. 4b-c, e-f) and within the undamaged parts of the crystals (Fig. 4d). Poorly preserved specimens including fragments and intact platelets with the central area missing are preferentially found in the cavities (Figs. 4b-c, f). However, the preservation of fine details on the platelets allows species-level 3 identification. Reticulofenestra minuta is by far the most abundant taxon forming dominant or monospecific assemblages (Figs. 4bd). Rare specimens of Reticulofenestra haqii, Reticulofenestra antarctica, Calcidiscus leptoporus, Coccolithus pelagicus, Helicosphaera carteri, Sphenolithus abies, Umbilicosphaera jafari and Chiasmolithus sp. are also observed (Fig. 4f). All taxa with the exception of Chiasmolithus sp., which went extinct in the Chattian, have chronostratigraphic ranges that overlap with Stage 3.

Sr/ 86 Sr results

Gypsum samples are characterized by a narrow range of 87 Sr/ 86 Sr (0.708704-0.708813; Table 1, Fig. 2). These data are substan-tially lower than 87 Sr/ 86 Sr ratios of coeval seawater (∼0.709022; McArthur et al., 2012) and consistent with published 87 Sr/ 86 Sr ratios from Eraclea Minoa (García-Veigas et al., 2018) and from selenites of the Siculiana Marina section (0.708700-0.708760; Keogh and Butler, 1999), ∼20 km to the SE (Fig. 1). Gypsum data from individual gypsum beds are either constant (II) or they show small-scale variability, commonly a decline from bottom to top (I, III-VII). There is also a small, but analytically resolvable decrease in 87 Sr/ 86 Sr ratios from the bottom of the section (Gypsum I) to the top (Gypsum VII).

The 87 Sr/ 86 Sr range of sub-unit D's ostracods is substantially larger than the gypsum (0.708510-0.708827; Fig. 2). Four of these samples (three in cycle 3 and one in cycle 7) have 87 Sr/ 86 Sr ratios

Discussion

Coccoliths in gypsums

The presence of marine calcareous nannofossils within the Eraclea Minoa UG selenites indicates that they were settling out from 6 suspension while gypsum was growing on the basin floor. Fragmented, poorly preserved platelets with an empty central area are present, indicating that some degree of physical reworking took place. The single occurrence of Chiasmolithus sp. corroborates this. However, fragmented platelets are preferentially found in large, irregularly-shaped cavities (Figs. 4b-c). When pristine areas of the gypsum slices (Fig. 4d) or small cracks (Fig. 4e) are inspected, preservation is better and complete coccospheres are observed. These delicate structures do not typically withstand erosion and transport intact. Other observations that are difficult to reconcile with an entirely reworked origin include the scarcity of terrigenous particles, the absence of bioclasts other than coccoliths and the limited diversity of these assemblages, which are solely or largely dominated by R. minuta (Figs. 4b,. The absence of other marine organisms is indicative of a stressed environment in which only a few species could thrive. Interestingly, the dominant species, R. minuta, is regarded as an opportunistic taxon capable of inhabiting a wide range of salinity conditions, from brackish to hypersaline [START_REF] Auer | High-resolution calcareous nannoplankton paleoecology as a proxy for small-scale environmental changes in the Early Miocene[END_REF]. Taken together, these data suggest that while minor reworking occurred, at least R. minuta, which is abundant, thrived in the Caltanissetta Basin during selenite precipitation. Being that R. minuta is of a marine origin, its finding in the Stage 3 selenites of Eraclea Minoa indicates that the Caltanissetta Basin was, at least during the evaporitic phases, supplied with Atlantic seawater. The seawater-like δ 34 S SO4 signature of the gypsums is consistent with the nannofossil-based conclusion (Fig. 2; [START_REF] Manzi | A new facies model for the Upper Gypsum of Sicily (Italy): chronological and paleoenvironmental constraits for the Messinian salinity crisis in the Mediterranean[END_REF]García-Veigas et al., 2018).

Monospecific assemblages of dwarfed nannofossils, identified as likely belonging to the genus Reticulofenestra, have also been described from the thin calcareous laminae of the balatino lithofacies in ODP Site 653B (Tyrrhenian Basin; Pierre and Rouchy, 1990) and in the Montedoro section (Rouchy and Caruso, 2006), located some 50 km to the NE of Eraclea Minoa, and from the balatino lithofacies in ODP Site 968A (Eratosthenes Seamount; Blanc-Valleron et al., 1998), where again they are embedded within the gypsum crystals themselves. These findings in gypsums from intermediatedeep settings to the east and west of the Caltanissetta Basin suggest that the Stage 3 gypsum subbasins of the Mediterranean may have been, at least during the evaporitic phases, part of a single, Atlantic-fed water body (which we refer to as "Lago-Mare"; see next section) typified by stressful environmental conditions (Fig. 6). The Western Alborán, Adriatic and North Aegean basins, which are Stage 3 gypsum-free, are likely to have been sealed off from the gypsum-forming Lago-Mare water mass due to the presence of dividing sills and to have maintained, for at least part of Stage 3, different water masses subject to independent baselevel and hydrologic evolution (Fig. 6). Studies of the sedimentary and paleontological records of these basins support this conclusion (Pellen et al., 2017;Booth-Rea et al., 2018;Krijgsman et al., 2020).

87 Sr/ 86 Sr characterization of the "Lago-Mare" water mass

The gypsum 87 Sr/ 86 Sr data at Eraclea Minoa (0.708704-0.708813; Fig. 2), like the Stage 3 gypsums elsewhere in the Mediterranean (<0.708850; Roveri et al., 2014b), have substantially lower ratios than coeval ocean water (∼0.709022;McArthur et al., 2012). Modeled 87 Sr/ 86 Sr of mixtures of Atlantic seawater and circum-Mediterranean freshwater show that measurable deviations from the oceanic 87 Sr/ 86 Sr signature are achieved when river discharge constitutes at least 25% of the total water fluxes (Topper et al., 2014). The Upper Gypsum 87 Sr/ 86 Sr data therefore indicate that the Lago-Mare water mass was a marine water body strongly influenced by continental-derived sources with lower than marine 87 Sr/ 86 Sr.

If Mediterranean base-level at times of gypsum precipitation was low enough to expose basin margin sediments, rivers drain-ing these sediments may have influenced the water composition in Caltanissetta. Primary Lower Gypsum, both in situ (Stage 1) and reworked (Stage 2; Manzi et al., 2021), is abundant in the Caltanissetta Basin. Local rivers draining the margins may therefore have provided additional Ca 2+ and SO 2- 4 , but would have had relatively high 87 Sr/ 86 Sr (∼0.7090; Lugli et al., 2010). However, the catchments of these rivers today are dominated by Meso-Cenozoic limestones, which would have produced lower 87 Sr/ 86 Sr (<0.7080).

On a more regional scale, the presence of ostracods derived from Eastern Paratethys suggests that this water body was a contributor of low 87 Sr/ 86 Sr non-marine water (0.7084-0.7085;Grothe et al., 2020) to the Mediterranean during Stage 3 (e.g. Andreetto et al., 2021a). Climate model simulations for the late Miocene show that the hydrological budget of the Eastern Paratethys was positive during the MSC (Marzocchi et al., 2016). This means that this lowsalinity water flow to the Mediterranean was not restricted only to times of marl deposition, where the Paratethyan fauna is found, but was present also during the evaporitic phases. The Mediterranean's major rivers, including the Nile and Rhône (0.7060 and 0.7087 respectively; e.g. Topper et al., 2011), which today make up >80% of the total Mediterranean river discharge, are also likely to have contributed low 87 Sr/ 86 Sr water, along with North African fluvial systems that are currently dry (e.g. the Eosahabi River that flowed from Messinian Lake Chad), but for which there is evidence of late Miocene flow [START_REF] Griffin | The late Neogene Sahabi rivers of the Sahara and their climatic and environmental implications for the Chad basin[END_REF].

Mass-balance calculations were used to explore the relative proportions of Sr 2+ derived from the possible sources that explain the 87 Sr/ 86 Sr data at Eraclea Minoa (Fig. 5a). If we assume that water from Eastern Paratethys was characterized by river-like Sr concentrations (Figs. 5b-c), our model suggests that a maximum of ∼20% Atlantic seawater is required to reproduce the gypsum 87 Sr/ 86 Sr data (Fig. 5b). This result is in agreement with the box model calculations of Topper et al. (2014) and points to the presence of a strongly diluted marine water mass with a restricted Atlantic-Mediterranean connection during Stage 3 gypsum formation.

Restricted ocean inflow can be provided by an Atlantic-Mediterranean gateway either as a one-way flow or with a parallel outflow. The two-way flow configuration facilitates the formation of gypsum because it entails a larger inflow of water (and ions) and keeps the salinity below halite saturation (Krijgsman and Meijer, 2008). The presence of gypsum and absence of large volumes of demonstrably in-situ halite in Stage 1 and 3 units (Roveri et al., 2014a;Andreetto et al., 2021a) is consistent with this scenario. However, two-way exchange requires Mediterranean base-level to be equal with the Atlantic sea-level (Krijgsman and Meijer, 2008), which means the marginal Mediterranean subbasins remain permanently submerged. This is the case in Stage 1. However, the evidence of substantial base-level change (section 4.5) and the frequent switching from subaqueous to continental sedimentation on the margins (Andreetto et al., 2021a(Andreetto et al., , 2022) make persistent twoway exchange during the UG highly unlikely and support a small, unidirectional Atlantic input to the Mediterranean (Fig. 6). Existing box models simulations of the blocked outflow scenario result in rapid and voluminous halite precipitation (Krijgsman and Meijer, 2008). However, these simulations do not include the inflow of low-salinity water from Eastern Paratethys. One possibility is that inflow from Paratethys, with a small Atlantic contribution may have kept the Lago-Mare water mass below halite saturation, but allowed it, episodically, to achieve gypsum saturation.

Implications of UG cycles for base-level and insolation

Evaporite-marl alternations denote cycles of concentration and dilution of the water mass with respect to evaporitic ions and are commonly associated with water level changes (Lowenstein, River water was always draining into the Mediterranean with variable discharge at different moments during a precession/insolation cycle (Simon et al., 2017).

1988; Lugli et al., 2010). Reconstructing base-level changes in ancient evaporitic settings with confidence is difficult because reliable paleobathymetric indicators such as fossils and ichnofossils are commonly absent from high-salinity environments and evaporitic lithofacies are rarely diagnostic of any particular water depth.

The UG cycles at Eraclea Minoa are unusual because sub-unit D contains in-situ ostracods which do have paleobathymetric requirements. Living relatives of C. agrigentina (e.g. C . torosa), which dominates Biofacies 1 (Fig. 2), tolerate only a few tens of meters of water depth (Grossi et al., 2015;[START_REF] Meyer | Morphological and geochemical variations of Cyprideis (Ostracoda) from modern waters of the northern Neotropics[END_REF]. The more diverse Biofacies 2 at the top of sub-unit D, by contrast, is considered to indicate water depths up to 100 m (Stoica et al., 2016). Roveri et al., 2008). The question marks in the figure to the left refer to the uncertainty over the paleodepth during deposition of sub-units A, B and C. Note that our new interpretation of the UG challenges the long-lasting idea that PLG and UG reflect similar depositional mechanisms and paleoenvironmental conditions [START_REF] Manzi | A new facies model for the Upper Gypsum of Sicily (Italy): chronological and paleoenvironmental constraits for the Messinian salinity crisis in the Mediterranean[END_REF]Roveri et al., 2014a). This is consistent in a number of ways: (1) UG and PLG sequences comprise different facies (UG contains more abundant balatino gypsum and is missing the banded and branching selenite lithofacies of the PLG). ( 2) Euxinic shales interbedded with the PLG contain possibly in-situ marine fossils (e.g. Corbí and Soria, 2016), whereas the marls interbedded with the UG contain brackish and shallow water ostracods and brackish water foraminifera such as A. tepida (Fig. 2; Grossi et al., 2015). This is likely to reflect a greater proportion of marine water in the Mediterranean during Stage 1. ( 3) 87 Sr/ 86 Sr isotopes are consistent with the above interpretation, being, on average, close to, or within error of coeval ocean values on stage 1 gypsums (average of 0.708941; Lugli et al., 2010;Reghizzi et al., 2018) and substantially higher than those measured on Stage 3 gypsum samples (average of 0.708752; Fig. 7a). ( 4) The Stage 3 gypsum 87 Sr/ 86 Sr data are either constant within the layer or show a subtle decline from base to top (Figs. 5a,7). This is in contrast with a larger base to top increase in 87 Sr/ 86 Sr ratio in Stage 1 gypsum layers from the Vena del Gesso Basin (Northern Apennines; Reghizzi et al., 2018;Fig. 7a). ( 5) The PLG is largely restricted to marginal basins (e.g. Lugli et al., 2010;García-Veigas et al., 2018), whereas the UG is only found in intermediate-deep basins (Fig. 6). b: Summary outline of the main sedimentological (i), paleontological (ii) and geochemical (iii) data for each sub-units, their interpretation in terms of salinity (iv) and relationship with the base-level curve (v) and their paleoenvironmental and paleohydrological interpretation proposed in this work (vi).

The ostracod assemblages therefore suggest that the water depth increased during the deposition of sub-unit D (Fig. 7a).

The presence, in one marl layer interbedded with massive selenite beds (sub-unit C) in cycle 3, of both C. agrigentina and L. muelleri, which is extinct but also thought to have inhabited shallow water (Grossi et al., 2015), suggests that the selenites precipitated in shallow water. However, there is no physical or paleobiological evidence of very shallow water conditions in the cumulate gypsum of sub-unit B. In other ancient evaporitic settings, the transition from cumulates to bottom-grown selenites is commonly interpreted to indicate a shallowing-upward trend [START_REF] Ortí | Evaporite sedimentation in a tectonically active basin: the lacustrine Las Minas Gypsum unit (Late Tortonian, SE Spain)[END_REF]Warren, 2016) and sequences of cumulate and bottomgrown halite layers characterize lake level falls in the Dead Sea today [START_REF] Sirota | Sedimentology and stratigraphy of a modern halite sequence formed under Dead Sea level fall[END_REF]. It is possible therefore that the transition from sub-unit B to C equates to a shallowing-upward trend through gypsum deposition (Fig. 7a).

The distribution of gypsum and marl across the Caltanissetta Basin may also provide insights into the relative water depth of the two lithologies. Gypsum occurs throughout the basin with a fairly constant thickness, whereas the evaporite-free marl intervals thin eastward towards the foreland area and above intrabasinal structural highs [START_REF] Manzi | A new facies model for the Upper Gypsum of Sicily (Italy): chronological and paleoenvironmental constraits for the Messinian salinity crisis in the Mediterranean[END_REF], where they are barren or only 10 contain Biofacies 1 ostracods (e.g. Balza Soletta section; Maniscalco et al., 2019). [START_REF] Manzi | A new facies model for the Upper Gypsum of Sicily (Italy): chronological and paleoenvironmental constraits for the Messinian salinity crisis in the Mediterranean[END_REF]i n t e r p r e t e d this distribution as the consequence of continuous subaqueous deposition with eastward thinning of the marls related to increased distance from the terrigenous source. An alternative interpretation is that the baselevel was high enough to precipitate gypsum across the whole Caltanissetta Basin, but dropped during marl deposition, exposing more marginal areas to erosion and/or non-deposition and leaving only the deeper depocenters (e.g. Eraclea Minoa) underwater. During the subsequent base-level rise, recorded in Eraclea Minoa by the appearance of Biofacies 2 ostracods, these marginal areas were drowned again and accumulated thin, Biofacies 1-bearing marls. If correct, this interpretation suggests that there was a higher water level at the beginning of gypsum precipitation than at the beginning of sub-unit D marls (Fig. 7a).

If the Caltanissetta Basin was part of a larger Lago-Mare water mass, cyclical base-level changes seen there are likely to be the consequence of fluctuations in the discharge of one, or more, of the Mediterranean's water sources. The Mediterranean's freshwater budget is highly variable on precession timescales. It is enhanced during the insolation maxima, when more rainwater from North African monsoon and Atlantic storms is delivered to the Mediterranean (e.g. Marzocchi et al., 2016), and it reaches the lowest values during the insolation minima, when less freshwater is supplied and more water is lost by evaporation. It is therefore likely that base-level variations were in phase with the insolation forcing: the highstands coincided with insolation maxima and the lowstands with insolation minima (e.g. Simon et al., 2017). Previously, the gypsum beds at Eraclea Minoa have been correlated with increasing insolation following the insolation minima and the diatomites and ostracod-bearing marls tied to the transition from insolation maxima to minima (Fig. 7a; van der Laan et al., 2006;[START_REF] Manzi | A new facies model for the Upper Gypsum of Sicily (Italy): chronological and paleoenvironmental constraits for the Messinian salinity crisis in the Mediterranean[END_REF]. This is challenged by the relative base-level changes reconstructed here, which instead suggest that the base of the ostracod-bearing marls at Eraclea Minoa represents lowstand phases, corresponding to insolation minima (Fig. 7a), while the shallowing through gypsum sub-units B and C, corresponds to part of the transition from insolation maxima to minima (Fig. 7a).

This revised astronomical tuning dates the base of the Upper Gypsum at Eraclea Minoa to ∼5.515 Ma and the first indication of Paratethyan ostracods to ∼5.47 Ma. This is the oldest age known in the Mediterranean and contradicts the hypothesis that the migration of Paratethyan ostracods occurred as a single Mediterraneanwide event at 5.42 Ma (substages 3.1 and 3.2 (or Lago-Mare phase) of Roveri et al., 2014a; see also Roveri et al., 2008 andGrossi et al., 2011), possibly prompted by a sudden increase of the Paratethys outflow (Marzocchi et al., 2016).

Connectivity changes terminated gypsum precipitation

The UG sediments at Eraclea Minoa also contain evidence of changing palaeoenvironmental conditions. The vertical transition from sub-mm/mm-sized gypsum crystals in the balatino lithofacies (sub-unit B), which require supersaturated conditions favoring the formation of new nuclei over crystal growth (Lugli et al., 2010), to increasingly larger selenites (sub-unit C), which require gradually lower degree of brine supersaturation to grow larger (Lugli et al., 2010), is indicative of a progressive reduction of the brine concentration approaching the transition with sub-unit D. Then, the sub-unit C/D transition marks a major paleoenvironmental and paleohydrological event, testified by the abrupt cessation of gypsum precipitation and the appearance of shallow-water fauna. Microthermometry studies on fluid inclusions in gypsum lithofacies (Bigi et al., 2022) and thermodynamic calculations [START_REF] Aloisi | The geochemical riddle of "low-salinity gypsum" deposits[END_REF] show that the Eraclea gypsums precipitated from the evaporation of seawater at a salinity ∼ 115 g kg -1 . The ostracods in the overlying marls show a gradual transition from dominantly C . agrigentina (Biofacies 1), which tolerates a wide range of salinities (Grossi et al., 2015), to the more diverse Biofacies 2, which indicate fresher (oligohaline) conditions (Fig. 7b; Stoica et al., 2016).

One possibility is that the gradual drop in concentration with respect to gypsum and the salinity decrease across the sub-unit C/D transition and through sub-unit D were prompted by increased continental runoff [START_REF] Manzi | A new facies model for the Upper Gypsum of Sicily (Italy): chronological and paleoenvironmental constraits for the Messinian salinity crisis in the Mediterranean[END_REF]. In this case, however, a base-level rise should have accompanied the precipitation of gypsum, in order to bring enough ions for it to precipitate. In turn, this means that the base level would reach the highstand during deposition of Biofacies 2 (Fig. 7a; [START_REF] Manzi | A new facies model for the Upper Gypsum of Sicily (Italy): chronological and paleoenvironmental constraits for the Messinian salinity crisis in the Mediterranean[END_REF]. This scenario is in contradiction with our water level reconstruction, that indicate a shallowing through the gypsum and a deepening through the sub-unit D marls (Fig. 7a). The lessening of Atlantic inflow, which achieves both the base-level fall through sub-units B-C and the associated subtle declining trend in the 87 Sr/ 86 Sr ratios within individual gypsums (Figs. 2, 7a), seems to be a more plausible mechanism. In the context of deposition during a Mediterranean base-level fall, it is possible, therefore, that the abrupt transition from gypsum to ostracod-bearing marls was the consequence of the base level falling below the Caltanissetta Basin, leaving it as a perched lacustrine system above the deeper areas covered by the "Lago-Mare" water mass. The generally lower 87 Sr/ 86 Sr ratios and larger 87 Sr/ 86 Sr variations in the Eraclea Minoa marls, indicate a greater influence of freshwater inputs to the basin compared to the gypsums that have higher and less variable 87 Sr/ 86 Sr ratios. This suggests that the Caltanissetta Basin formed a perched lake during marl deposition at lowstands. Compared to seawater-influenced basins, lacustrine settings are, indeed, much more sensitive to subtle changes in fluvial discharge (e.g. [START_REF] Bista | Sr isotope-salinity modelling constraints on Quaternary Black Sea connectivity[END_REF], resulting in rapid changes in the 87 Sr/ 86 Sr ratio measured on the ostracods.

High amplitude fluctuations of the Mediterranean base-level

The same ostracods biofacies as those in sub-unit D are found both in marginal Mediterranean marls, where they typically alternate with continental facies, and in the deep basins (e.g. Sites 975B, 372, 129A; see Andreetto et al., 2021a), where the marls alternate with laminated and clastic gypsum (Andreetto et al., 2021a). For the deep basins to develop the shallow-water conditions required by the ostracods, the Mediterranean water level had to fall below the intermediate position of the Caltanissetta Basin. This means that Caltanissetta and deeper basins accumulated the shallow-water ostracod-bearing marls during lowstand phases of the Lago-Mare water, exposing the marginal basins to continental deposition and/or erosion (Fig. 8e). By contrast, higher Mediterranean water levels are required to populate the shallow marginal basins with these same Paratethyan ostracods (Stoica et al., 2016). This is supported by the 87 Sr/ 86 Sr ratios measured on some of these marginal basin ostracods, which point to mixing between high-87 Sr/ 86 Sr local fluvial waters with the Lago-Mare water mass (Andreetto et al., 2021b(Andreetto et al., , 2022)). During these highstand phases of marginal basin ostracod-rich marl deposition, the Caltanissetta and deep basins were experiencing much deeper water conditions (i.e. several hundreds of meters; Fig. 8b).

The depth constraints provided by the shallow-water ostracod assemblages indicate that marl deposition was not synchronous across the Mediterranean during base-level highstands at insolation maxima (Roveri et al., 2008(Roveri et al., , 2014a)), but sequential, tracking Mediterranean-wide, high-amplitude (several hundred to thousands of meters) base-level fluctuations . This model of high-amplitude base level changes, earlier hypothesized by Hardie and Lowenstein (2004) and Rouchy and Caruso (2006) times both relatively full (Fig. 8b) and empty (Fig. 8e) during each Stage 3 precession cycle. This model helps reconcile observations that support the full versus desiccated debate that has afflicted the MSC Stage 3 successions for decades (Andreetto et al., 2021a).

Implications of Atlantic-Mediterranean connectivity for ocean chemistry

On long timescales (>100 kyr), marine alkalinity and CaCO 3 burial are balanced by the response of silicate weathering to changes in atmospheric CO 2 and associated temperature. Many biogeochemical models (e.g. [START_REF] Berner | A model for atmospheric CO sub 2 over Phanerozoic time[END_REF][START_REF] Lenton | COPSE reloaded: an improved model of biogeochemical cycling over Phanerozoic time[END_REF] assume that on these long timescales perturbations to seawater chemistry caused by evaporite precipitation are balanced by evaporite dissolution during weathering. However, evaporites form and dissolve three orders of magnitude more quickly than other sediments (Nichols et al., 2009) and this approach fails to consider whether abrupt and short-term evaporite-driven chemical perturbations can propagate into aspects of the carbon cycle that have longer-term climate implications. The box model study by Shields and Mills (2020) explores the impact of gypsum on ocean chemistry and its knock-on consequences for CaCO 3 burial, atmospheric CO 2 and global temperature by allowing the non-steadystate ocean calcium inventory to be driven by CaSO 4 precipitation and the return flux to the ocean from weathering. Their results demonstrate that a decrease in ocean [Ca] as a consequence of gypsum deposition, because it is not associated with an alkalinity change, leads to decreasing CaCO 3 burial and higher oceanic pH, resulting in decreasing atmospheric CO 2 and global cooling (Shields and Mills, 2020). However, their sensitivity experiments impose a constant gypsum precipitation and weathering forcing for 1 million years which is not a realistic reflection of the pattern of ocean [Ca] extraction as a consequence of the MSC. Here we explore what a realistic evaporite precipitation and weathering flux would look like over the duration of the MSC and consider its possible implications.

In contrast with the constant precipitation driver modeled by Shields and Mills (2020), the precessional gypsum-marl cyclicity that characterizes both MSC Stage 1 and 3 suggests that extraction of Ca 2+ and SO 2- 4 ions from the global ocean was episodic and probably endured for only a few thousand years in each precession cycle. Further modeling experiments are required to evaluate the degree to which the regular switching on and off of gypsum precipitation modifies the modeled consequences for carbonate burial, atmospheric CO 2 and global temperatures (Shields and Mills, 2020) as a result of (partial) recovery of the system between evaporite pulses. However, without near synchronous return of Ca 2+ ions to the ocean as a result of evaporite weathering, the reduction in oceanic [Ca] is unlikely to have bounced back completely, leading to a progressive decline in seawater [Ca] from one precessional evaporite event to the next. Stage 2 evaporites are heavily dominated by halite, not gypsum. However, it is possible that constant extraction of these Ca 2+ ions did characterize the ∼50 kyr duration of this phase of the MSC, since the volumetrics of the halite succession preserved, requires constant input of ocean water (Krijgsman and Meijer, 2008) which must also have contained Ca 2+ and SO 2- 4 . Taken together then, the rate of Ca 2+ ion extraction over the course of the entire MSC is likely to have varied and to have been most substantial during Stage 2.

Recent evaporite volume estimates for the MSC (Haq et al., 2020) are broadly similar to those used by Shields and Mills (2020) to estimate the scale of the Ca 2+ extraction during salt giant formation. Both are based on the volumes of evaporites that are still preserved in the Mediterranean today, and therefore represent a minimum. That being the case, whether pulsed or continuous, this Late Miocene episode of CaSO 4 extraction may have contributed to the cooling seen in time-equivalent SST records (Herbert et al., 2016).

Although the net effect of the MSC was to extract Ca 2+ and SO 2- 4 from the global ocean, the episodic nature of gypsum accumulation and Mediterranean-Atlantic connectivity also implies that the Mediterranean was, intermittently and at time-scales shorter than the MSC, a net source of Ca 2+ and SO 2- 4 to the Atlantic. Evaluating episodic net Ca 2+ and SO 2- 4 fluxes to the ocean during the MSC is particularly challenging since it depends not on the gypsum preserved today in the Mediterranean, but rather on what was originally precipitated and has been already dissolved (Wortmann and Payton, 2012). However, what we can evaluate from the Mediterranean's geological record is the nature of its connectivity with the Atlantic and from this it is possible to deduce the timing and likely duration of episodic net exchange of Ca 2+ and SO 2- 4 between the Mediterranean to the global ocean.

Evidence from both 87 Sr/ 86 Sr data and box-modeling indicate that Stage 1 gypsum was precipitated while the Mediterranean experienced restricted, but two-way exchange with the Atlantic under conditions which broadly maintained eustatic sea level in the Mediterranean (e.g. Flecker et al., 2015 and references therein). This configuration will have resulted in net extraction of Ca 2+ and SO 2- 4 ions from the ocean, at least during gypsum precipitation. Maximum dissolution of gypsum occurs during subaerial exposure with the formation of aquifers associated with active hydrology [START_REF] Gustavson | Evaporite dissolution and development of karst features on the Rolling Plains of the Texas Panhandle[END_REF]. Since there is no evidence of substantial base-level fall during the marl intervals in Stage 1, it is probable that there was relatively little weathering of the gypsum and therefore a limited return flux of these ions to the ocean.

The connectivity and base-level configurations during stages 2 and 3 are substantially different. Outflow to the Atlantic ceased, base-level fell and then, as shown in this study, fluctuated substantially during Stage 3. This exposed the Stage 1 gypsum that precipitated on the margins and stage 2 and 3 gypsum reworked or precipitated in deeper settings of the Mediterranean, to dissolution. However, without Mediterranean outflow, these weathered evaporitic ions were trapped in the Mediterranean and had no influence on global ocean chemistry. Outflow to the Atlantic started abruptly at the beginning of the Pliocene with the termination of the MSC, triggering the return of two-way exchange and the reestablishment of open marine conditions in the Mediterranean.

In summary, the MSC evaporite weathering flux to the ocean was episodic and largely out of phase with gypsum precipitationdriven extraction, potentially resulting in dynamic changes to ocean chemistry on the timescale of the MSC. Interestingly, the main impact Shields and Mills (2020)f o u n d of modeling a 1 million year CaSO 4 weathering event was a rapid rise in CaCO 3 burial and a transient peak in atmospheric CO 2 concentration and associated warming, that occurred only 150 kyrs after initiation. With this in mind, it is tempting to attribute at least part of the abrupt warming in early Pliocene sea surface temperature records (Herbert et al., 2016) to a Mediterranean-derived pulse of Ca 2+ ions. Testing this hypothesis will require more sophisticated biogeochemical modeling and better quantitative constraints on the chemistry of Mediterranean outflow. If true however, it would suggest that the most important impact of the Zanclean deluge was not the sudden inundation of a Mediterranean empty of water for hundreds of thousands of years, but the abrupt flushing into the Atlantic of the ponded brine and weathering products of a salt giant.

Conclusions

The UG selenites at Eraclea Minoa contain in-situ marine calcareous nannofossils and display 87 Sr/ 86 Sr ratios which indicate that a small Atlantic flux (≤20%) supplied the Mediterranean with Ca 2+ and SO 2- 4 ions during gypsum precipitation. Mass-balance calculations suggest that the remaining ≥80% of the Lago-Mare water mass was supplied by Eastern Paratethys and large Mediterranean rivers. Gypsum precipitation, first from settling crystals nucleated in the water column (balatino, sub-unit B) and then from bottom-grown selenites (sub-unit C), occurred during a phase of base-level fall coincident with the transition from insolation maxima to insolation minima. During insolation minima, low diversity ostracod-bearing marls (Biofacies 1 of sub-unit D), that constrain the water depth to a few tens of meters, were deposited in an isolated, lacustrine Caltanissetta Basin perched above the Lago-Mare water mass, which persisted in the deep subbasins. The transition from Biofacies 1 to 2 indicates rising water levels in the lake. Similar ostracod-bearing marls are found in both shallower and deeper settings across the Mediterranean, indicating that marl deposition was asynchronous during Stage 3, tracking high-amplitude, Mediterranean-wide base level changes (Fig. 8a). Sedimentological, geochemical and biological data from Eraclea Minoa suggest that the Mediterranean experienced (at least) 7 substantial, cyclic base-level changes during Stage 3. A revised astronomical tuning indicates that MSC Stage 3 started at 5.515 Ma, with the first occurrence of Paratethyan ostracods on Sicily at 5.47 Ma. This model for (upper) gypsum formation, which envisages periods during which the Mediterranean was both relatively full and empty, helps to reconcile many of the discrepancies between offshore-and onshore-based interpretations of the final phase of the Messinian Salinity Crisis. It also suggests that any impact of the Mediterranean salt giant formation and evaporite weathering during stages 13

Introduction

At the end of the Miocene, starting from about 7.2 Ma, the progressive reduction of water exchanges between the Atlantic Ocean and the Mediterranean Sea dramatically affected Mediterranean environments, increasing their sensitivity to orbitally-driven climate perturbations (Bulian et al., 2021;Hsü et al., 1973;[START_REF] Kouwenhoven | Paleoenvironmental evolution of the eastern Mediterranean during the Messinian: constraints from integrated microfossil data of the Pissouri Basin (Cyprus)[END_REF]. Hydrological restriction of the Mediterranean led to an increase of salinity, a decrease in water column oxygenation and changes in biotic assemblages (Bulian et al., 2021;Cita et al., 1978;[START_REF] Kouwenhoven | Paleoenvironmental evolution of the eastern Mediterranean during the Messinian: constraints from integrated microfossil data of the Pissouri Basin (Cyprus)[END_REF]. Thereafter, the sensitivity of the Mediterranean to precessional cycles became significantly more pronounced, and in turn greatly impacted the biogeochemistry of the basins (Natalicchio et al., 2019;[START_REF] Sabino | The response of water column and sedimentary environments to the advent of the Messinian salinity crisis: insights from an onshore deep-water section (Govone, NW Italy)[END_REF]. This has been recorded as cyclical deposition of sapropel/diatomite-rich marls (Krijgsman et al., 1999) and marl/carbonate alternations [START_REF] Mancini | Calcareous nannofossil size and abundance response to the Messinian salinity crisis onset and paleoenvironmental dynamics[END_REF] corresponding to periods of summer insolation maxima and minima, respectively. Lithological cyclicity is accompanied by variations in the size and abundance of fossils of marine fauna [START_REF] Mancini | Calcareous nannofossil size and abundance response to the Messinian salinity crisis onset and paleoenvironmental dynamics[END_REF], as well as oscillations in the composition and isotopic signature of lipid biomarkers (Natalicchio et al., 2019;[START_REF] Sabino | Climatic and hydrologic variability in the northern Mediterranean across the onset of the Messinian salinity crisis[END_REF]. Together, this evidence suggests the existence of a precessional-scale alternation between * Corresponding author. E-mail address: guibourdenche@ipgp.fr (L. Guibourdenche). a highly productive/stratified water column and a vertically mixed water column.

The progressive intensification of hydrological restrictions between the Atlantic and the Mediterranean culminated at 5.97 Ma with the onset of evaporite precipitation that marked the beginning of the Messinian Salinity Crisis (MSC). During the first phase of the MSC, up to 220 m of cyclical alternations of gypsum and organic-rich deposits (characteristic of arid and humid periods, respectively) forming the Primary Lower Gypsum (PLG) unit (Fig. 1) accumulated in the shallow part of marginal to intermediate depth basins of the Mediterranean (Lugli et al., 2010;Raad et al., 2021;Roveri et al., 2014). During this time, marginal and intermediate depth basins still received significant marine water inputs, as indicated by the 87 Sr/ 86 Sr and δ 34 S of gypsum and by the occurrence of marine biota (Lugli et al., 2007, 2010 andreferences therein;Pellegrino et al., 2021;Reghizzi et al., 2018).

During humid periods of precession cycles, non-evaporitic interbeds deposited in stratified basins where N 2 -fixing bacteria significantly contributed to primary productivity (Isaji et al., 2019) and anoxygenic phototrophic bacteria populated the upper water column (Sinninghe Damsté et al., 1995). Meanwhile, microbial mats dominated by colorless sulfide oxidizing bacteria (CSB) thrived at the seafloor (Dela Pierre et al., 2014). Anoxygenic phototrophic bacteria and CSB mats develop in modern euphotic zone and benthic boundary respectively. They reoxidize H 2 S fluxes produced by microbial sulfate reduction (MSR) back to elemental sulfur or sulfate [START_REF] Brüchert | Biogeochemical and Physical Control on Shelf Anoxia and Water Column Hydrogen Sulphide in the Benguela Coastal Upwelling System off Namibia[END_REF]Callbeck et al., 2021; https://doi.org/10.1016/j.epsl.2022.117676 0012-821X/© 2022 Elsevier B.V. All rights reserved. Jørgensen, 2021). The presence of sulfide oxidizing metabolisms in both the water column and at the sediment-water interface strongly suggest the prevalence of euxinic conditions during marls/organic-rich shale deposition.

Precession maxima (arid periods) were characterized by rapid vertical accretion of bottom grown gypsum crystals (Reghizzi et al., 2018) -which allowed for the relatively good preservation of organic remains (Dela Pierre et al., 2015;Schopf et al., 2012) (Fig. 2). These bottom-grown gypsum crystals make up the bulk of the PLG unit, and contain abundant planktic marine diatoms, implying enhanced primary productivity in the upper water column during gypsum deposition (Dela Pierre et al., 2015;Pellegrino et al., 2021;Schopf et al., 2012).

The molecular fossil assemblages entrapped within Messinian gypsum crystals and their compound specific carbon isotopes are indicative of a diverse microbial community, including marine Thaumarchaeota communities living in the upper water column, and sulfate-reducing bacteria and CSB-forming microbial mats at the sea floor (Natalicchio et al., 2022). The latter benthic microbial mats are fossilized within the bottom-grown gypsum in the form of densely-packed networks of filamentous structures ("spaghettilike" structures) together with pyrite and dolomite, implying that MSR activity occurred during gypsum formation (Dela Pierre et al., 2015;Natalicchio et al., 2022;[START_REF] Panieri | Microbial communities in Messinian evaporite deposits of the Vena del Gesso (northern Apennines, Italy)[END_REF]Schopf et al., 2012) (Fig. 2D-I).

Together, these petrographic and geochemical observations point towards a highly dynamic sulfur cycle during PLG formation, with sulfur cycled amongst microbial sulfate reducers and sulfideoxidizing bacteria, as well as precipitated in its oxidized form as gypsum (CaSO 4 .2H 2 O).

Marine carbon and sulfur cycles are tightly linked, with MSR driving the remineralization of between 12 and 65% of the organic carbon flux to the seafloor [START_REF] Bowles | Global rates of marine sulfate reduction and implications for sub-sea-floor metabolic activities[END_REF][START_REF] Bradley | Widespread energy limitation to life in global subseafloor sediments[END_REF]Canfield et al., 2005). Constraining sulfur cycling during the formation of giant sulfate deposits is therefore critical to understand their significance for the carbon cycle. The accumulation of large gypsum deposits -comparable to the Mediterranean PLG-are known to greatly affect the long term carbon cycle and atmospheric oxygen levels by the removal of substantial amount of Ca 2+ and SO 4 2-ions from the ocean, (e.g. Halevy et al., 2012;Shields and Mills, 2020). However, the role of associated MSR and sulfide oxidation on carbon and sulfur cycles of basins where gypsum formed has been overlooked. This complex sulfur cycling during PLG formation is likely to have been widespread across Mediterranean marginal basins (Natalicchio et al., 2022;Pellegrino et al., 2021)) and raises fundamental questions about the biogeochemical dynamic of these basins, including for instance: (i) What was the relative importance of microbially driven reduction of sulfate and reoxidation of sulfide, compared to the burial of sulfate through gypsum precipitation? (ii) How did sulfur cycling impact carbon cycling in Messinian Mediterranean marginal basins? (iii) How does this ancient system compare with modern marine environments? In this study, we investigated the most complete PLG succession of the Mediterranean area in the Vena del Gesso basin (VdG) (Northern Apennines, Italy) (Fig. 1b). MSR and re-oxidative sulfur cycling result in specific multiple sulfur isotopes signatures, providing a tool to disentangle the relative influence of concurrent sulfur redox processes [START_REF] Johnston | Sulphur isotopes and the search for life: strategies for identifying sulphur metabolisms in the rock record and beyond[END_REF]Ono et al., 2006). Therefore, to constrain the sulfur biogeochemical cycle during PLG deposition, we carried out measurements of multiple sulfur isotope ratios ( 34 S/ 32 S and 33 S/ 32 S) on gypsum sulfate and reduced sulfur minerals which co-precipitated with gypsum or that we extracted from marls, coupled to the measurement of oxygen isotopes of gypsum sulfate ( 18 O/ 16 O). This approach enabled us to estimate the relative fluxes of sulfide and sulfate in the basin and in turn to estimate the gross flux of organic carbon involved in MSR activity. In modern settings, this is typically achieved with incubation experiments and/or analyses of pore water geochemical profilestools that are not available when studying the rock record.

Geological setting

The PLG was deposited during the first stage of the MSC in shallow to intermediate marginal basins of the Mediterranean, including southern Spain, Italy, Hellenic arc and Cyprus arc (Krijgsman et al., 1999;Lugli et al., 2010 and references therein) (Fig. 1c). This unit was deposited in approximately 380 kyrs (Roveri et al., 2014) and is formed by lithological cycles consisting of organic-rich (marls, shale and carbonate)/gypsum couplets (Fig. 1a). The deposition of these interbeds was controlled by precession-driven climate oscillations (Lugli et al., 2010;Roveri et al., 2014). Tuning of the lithological cycles to the astronomical reference curve (Fig. 1a) allows the precise dating of each cycle and bed by bed correlation of the PLG unit at the Mediterranean scale (see the pink dots in Fig. 1c) (Krijgsman et al., 1999;Roveri et al., 2014). Here we studied the VdG in the Northern Apennines (Italy) that exposes the most complete succession of PLG deposits in the whole Mediterranean area (Fig. 1b). Each PLG cycle of the VdG basin starts with a thin (dm to m thick) organic matter-rich (total organic carbon from 0.2 to 3.1%; Lugli et al., 2007) marly bed (Fig. 2a).

The thin marly beds are followed in each PLG cycle by a 4 to 30 m thick bed of primary gypsum (Fig. 2a). The gypsum beds of the first three cycles (PLG1 to PLG3) are mostly composed of bottomgrown, vertically oriented, twinned selenite crystals (up to 2.5 m high) of the massive selenite lithofacies (Lugli et al., 2010).

Starting from the gypsum beds of the 4th PLG cycle, the massive selenite lithofacies is followed upwards by the banded selenite lithofacies composed of less than 10 cm-thick crusts of bottomgrown twinned gypsum crystals separated by thin carbonate laminae possibly reflecting cyclical fluctuation of the gypsum saturation interface (Lugli et al., 2010) (Fig. 2b).

From the 6 th PLG cycle to the top of the section, massive and banded selenite are followed by the branching selenite lithofacies consisting of cm-sized clear bottom-grown selenite crystals growing inclined or horizontally and grouped together in nodules or lenses separated by thin fine grained carbonate or gypsum laminae (Lugli et al., 2010). The absence of karst features in any of these lithofacies suggests that PLG gypsum are fully subaqueous deposits formed in a basin permanently covered by waters supersaturated with respect to gypsum. The sulfur and oxygen stable isotope compositions of pore water dissolved sulfate are frequently used to investigate the biogeo- δ 34 S SO4 and δ 18 O SO4 isotopic composition of the dissolved sulfate ion in equilibrium with the PLG unit from VdG basin (dark pink) and the PLG unit of other marginal basins (light pink) (see Supplementary section 2.1 for references and 2.4 for calibration of δ 34 S SO4 and δ 18 O SO4 isotopic data from this study and from the literature).

Isotopes to decipher sulfur cycling

The range of isotope compositions of sulfate-bearing minerals and of sedimentary sulfides are represented as grey and light brown boxes, respectively (see Supplementary section 2.1 for references). The possible range of isotopic composition of riverine inputs is bracketed by the isotopic composition of sulfate-bearing minerals and sedimentary sulfides (see Supplementary section 2.1 for further details on riverine fluxes). The possible evolution of the isotopic composition of residual sulfate during MSR is indicated by the orange shaded area. The isotope composition of sulfate dissolved in top few centimeters of modern marine environments characterized by near complete reoxidation of H 2 S is represented by the blue box (Aller et al., 2010). (For interpretation of the colors in the figure, the reader is referred to the web version of this article.) chemical sulfur cycling in modern marine sediments (e.g. Aller et al., 2010;Blonder et al., 2017;[START_REF] Gilhooly | A comprehensive sulfur and oxygen isotope study of sulfur cycling in a shallow, hyper-euxinic meromictic lake[END_REF]. These stable isotopic compositions are expressed in delta (δ) notations, representing the deviation of 34 S/ 32 S or 18 O/ 16 O ratios from international VCDT and SMOW standards respectively:

δ( ) = R measured R standard -1 × 1000 (1)
With R measured corresponding to the measured 34 S/ 32 S or 18 O/ 16 O ratios and R standard corresponding to the VCDT 34 S/ 32 S ratio or SMOW 18 O/ 16 O ratio.

Sulfate-reducing microbes preferentially metabolize 32 S-and 16 O-bearing dissolved sulfate. Therefore, as dissolved sulfate concentration decreases and hydrogen sulfide concentration increases, residual pore-water sulfate is progressively enriched in 34 S and 18 O (Fig. 3) (Brunner et al., 2005).

However, the effect of MSR on sulfate and hydrogen sulfide concentration profiles and isotopic compositions can be obscured by re-oxidative processes; this happens because 34 S-depleted sulfur metabolized by MSR is re-injected back into the 34 S-enriched sulfate pool with little or no fractionation (Fig. 3). This results in a smaller net enrichment of dissolved sulfate δ 34 S due to MSR, and a lower net production of sulfide/consumption of sulfate. Importantly, if all the hydrogen sulfide produced by MSR is entirely re-oxidized back to sulfate (Fig. 3) (Aller et al., 2010;Blonder et al., 2017;[START_REF] Brüchert | Biogeochemical and Physical Control on Shelf Anoxia and Water Column Hydrogen Sulphide in the Benguela Coastal Upwelling System off Namibia[END_REF]Callbeck et al., 2021), the sulfur cycle is "cryptic" -since there is neither change in sulfur species concentrations nor net isotopic effect on the δ 34 S isotopic composition. Furthermore, the shift in δ 18 O sulfate isotopic values is often so small (<4 ) (Aller et al., 2010;Blonder et al., 2017) that it can be difficult to recognize the effect of cryptic sulfur cycling from natural variability of oceanic sulfate isotopic composition using this isotope system (Fig. 3) (Markovic et al., 2016).

Multiple sulfur isotopes

Multiple sulfur isotopes (including 32 S, 33 S, 34 S) have the potential to distinguish biogeochemically distinct hydrogen sulfide re-oxidation pathways occurring concomitantly with MSR -the signal of which might otherwise be obscured using traditional isotope systems [START_REF] Johnston | Sulphur isotopes and the search for life: strategies for identifying sulphur metabolisms in the rock record and beyond[END_REF]Ono et al., 2006;[START_REF] Zhang | Redox chemistry changes in the Panthalassic Ocean linked to the end-Permian mass extinction and delayed Early Triassic biotic recovery[END_REF].

The partitioning of 34 S ( 34 S/ 32 S) relative to 33 S ( 33 S/ 32 S) is predictable under thermodynamic equilibrium, because it depends on the relative mass of the isotope considered. However, intracellular biochemical branching processes can shift the minor 33 S sulfur isotope partitioning ( 33 S/ 32 S) thus resulting in a deviation from the expected thermodynamic equilibrium relationship (Ono et al., 2006). This deviation is expressed as 33 S where:

33 S i = 33 R measured 33 R VCDT -1 × 1000 - 34 R measured 34 R VCDT -1 λ eq × 1000 (2)
The measurement of δ 34 S-33 S thus provides additional information as deviation from thermodynamic equilibrium can be linked to biological fractionation processes and mixing proportions of end-member sulfur species involved in sulfur cycling (Fig. 4) (Ono et al., 2006).

The deviation from thermodynamic equilibrium is also expressed by the 33 λ notation that defines the relationship between isotopic fractionation for 33 S and 34 S isotopes among two sulfur species:

33 λ j = ln( 33 α j ) ln( 34 α j )

with λ for thermodynamic equilibrium 33 λ eq = 0.515, and x α j =

x/32 R product / x/32 R reactant , and x = 33 or 34.

Combining the 33 λ and the classical 34 ε [= ( 34 α j -1) × 1000] notations is particularly useful as λ and ε fractionation factors are dependent solely on the biogeochemical processes in question (e.g. MSR, biotic or abiotic sulfide oxidation). In addition, the combination of 33 λ and 34 ε is not sensitive to the absolute isotopic value of the reacting sulfur species (Fig. 5) or the almost invariant temperatures relevant for sedimentary environments [START_REF] Eldridge | Theoretical estimates of equilibrium sulfur isotope effects in aqueous sulfur systems: highlighting the role of isomers in the sulfite and sulfoxylate systems[END_REF].

Material and methods

A total of 88 samples that are representative of each lithofacies of all cycles from the VdG succession were analyzed (Fig. 1A). Gypsum and marls from PLG cycle 1-2 were sampled in the Monticino quarry (44 • 13 ′ 29 ′′ N, 11 • 45 ′ 43 ′′ E) whereas cycles 3 to 16 were sampled in the Monte Tondo quarry (44 • 15 ′ 04 ′′ N, 11 • 40 ′ 13 ′′ E). Care was taken to avoid sampling the weathered surface and visible fractures in order to limit contamination.

Petrographic analyses

11 oriented samples were taken from the various gypsum lithofacies for petrographic observations. Fifteen petrographic thin sections and 10 thin slices (∼1 mm thick) were produced from the most representative samples by cleaving the selenite crystals along the main (010) cleavage planes with a razor blade. These were studied using transmitted, reflected, and UV-visible light microscopy, using a Leika DM 2700 P microscope with excitation filters 340-380 Nm at the Department of Earth Sciences of the University of Torino.

Oxygen and multiple sulfur isotope analyses

Sulfate from gypsum and iron sulfide from gypsum and marls were first reduced and transformed into Ag 2 S by wet chemistry. The multiple sulfur isotope composition was then measured using a MAT-253 mass spectrometer. Oxygen isotopic composition of sulfate from gypsum were measured on samples precipitated as BaSO 4 (the full methods with references are detailed in supplementary material section 1).

Geochemical modeling 4.3.1. Description of the model

We interpret the multiple S isotope data using a steady-state, 3-box model based on a chemical and isotopic mass balance of sulfur in the SO 4 2-, S • and H 2 S forms. In the model (see Fig. 6) sulfate is provided by oceanic (F in ) and riverine inputs (F riv ) whereas it is removed from the basin through outflow into the open ocean (F out ) or gypsum precipitation (F gyp ). In addition, the sulfate can be reduced to hydrogen sulfide by microbial sulfate reduction (F msr ), and hydrogen sulfide can be either buried as solid sulfide (F is ) or reoxidized (F ox ). Sulfide reoxidation (biotic or abiotic) produces sulfate or an intermediate sulfur compound (elemental sulfur or sulfite) (F ox ). Finally, intermediate sulfur compounds can be reoxidized back to sulfate (F ox ) or further disproportionated into hydrogen sulfide (F dh ) and sulfate (F do ) (Fig. 6).

The model considers the sulfur isotopes 32 S, 33 S and 34 S and calculates the sulfur isotope composition (δ 34 S, δ 33 S, 33 S) of gypsum, S • and solid sulfide formed in the marginal basin. Model equations consist of two mass balances for each of the three boxes: one chemical mass balance for total sulfur and one isotopic mass balance for the heavy sulfur isotopes, representing 33 S or 34 S (see supplementary section 5 for a full model description).

Fractionation factors

Sulfur redox or phase changes are often associated with a differential partitioning of sulfur isotopes. Therefore, the expressed fractionation ε net and λ net between two co-occurring sulfur species of interest (typically sulfate and sulfide) results from the net effect of multiple processes involved in the cycling of these sulfur species (Fig. 5). To model the biogeochemical processes at play during the genesis of our samples, we assigned fixed fractionation factors established in pure culture experiments or environmental studies to each of the fractionating process considered in our model (see supplementary section 4.1), with the exception of sulfate reduction that potentially expresses a large range of isotopic fractionation [START_REF] Gomes | Sulfur isotope fractionation in modern euxinic systems: implications for paleoenvironmental reconstructions of paired sulfatesulfide isotope records[END_REF]Sim et al., 2011b;[START_REF] Wing | Intracellular metabolite levels shape sulfur isotope fractionation during microbial sulfate respiration[END_REF]. In the following paragraph, we detail the approach we chose to simulate fractionation of multiple sulfur isotopes during MSR.

Laboratory experiments and studies of natural environments suggest that there is a first-order control of the specific cellular rate of MSR (csSRR) on the extent of sulfur isotopic fractionation during MSR, with fractionation approaching equilibrium values (i.e. ε eq =-70 and λ eq = 0.515) when csSRR reaches minimum values (Sim et al., 2011a,b;[START_REF] Stam | 34S/32S fractionation by sulfate-reducing microbial communities in estuarine sediments[END_REF] (Fig. 5). The laboratoryobserved csSRR at which these relationships have been established are often 1-3 orders of magnitude higher than the csSRR usually reported from natural marine environments [START_REF] Colangelo-Lillis | Diversity decoupled from sulfur isotope fractionation in a sulfatereducing microbial community[END_REF]Jørgensen, 2021 and references therein). Some authors therefore argue that csSRR in marine environments are so low that cells always express fractionation factors close to equilibrium values (ε MSR ≈-70 and λ MSR ≈0.515) (e.g. [START_REF] Masterson | Interpreting multiple sulfur isotope signals in modern anoxic sediments using a full diagenetic model (California-Mexico margin: Alfonso Basin)[END_REF]. These authors propose that when apparent fractionation factors between sulfate and sulfide (ε net and λ net ) deviate from equilibrium values, the deviation is mainly a consequence of sulfide pooling produced by sedimentary or early diagenetic processes (see section 5.2.2.1 for further explanations on this mechanism) [START_REF] Masterson | Interpreting multiple sulfur isotope signals in modern anoxic sediments using a full diagenetic model (California-Mexico margin: Alfonso Basin)[END_REF].

ε net > -70 and λ net < 0.515 are often observed in systems where depositional controls are absent or negligible, implying that natural populations of microbial sulfate reducers can produce a large range of isotopic fractionations [START_REF] Colangelo-Lillis | Diversity decoupled from sulfur isotope fractionation in a sulfatereducing microbial community[END_REF][START_REF] Gilhooly | A comprehensive sulfur and oxygen isotope study of sulfur cycling in a shallow, hyper-euxinic meromictic lake[END_REF]Gomes and Hurtgen, 2015 and references therein;literature review in Sim et al., 2011a). In addition, modeling studies also suggest a first order control of csSRR on fractionation during MSR [START_REF] Bradley | Patterns of sulfur isotope fractionation during microbial sulfate reduction[END_REF][START_REF] Wing | Intracellular metabolite levels shape sulfur isotope fractionation during microbial sulfate respiration[END_REF]. Howbeit, secondary factors independent from csSRR might also influence the extent of isotopic fractionation expressed during MSR (e.g. [START_REF] Bradley | Patterns of sulfur isotope fractionation during microbial sulfate reduction[END_REF][START_REF] Brüchert | Controls on stable sulfur isotope fractionation during bacterial sulfate reduction in Arctic sediments[END_REF][START_REF] Colangelo-Lillis | Diversity decoupled from sulfur isotope fractionation in a sulfatereducing microbial community[END_REF][START_REF] Stam | 34S/32S fractionation by sulfate-reducing microbial communities in estuarine sediments[END_REF][START_REF] Wing | Intracellular metabolite levels shape sulfur isotope fractionation during microbial sulfate respiration[END_REF]. Although it is not within the scope of this work to reconcile laboratory experiments with field observations and modeling studies, we point out that the correct estimation of csSRR in natural environments (in mol.cell -1 .day -1 ) relies on: (i) a precise estimate of the number of physiologically active sulfate reducing microbial cells, and (ii) the absence of significative re-oxidative sulfur cycling. Any overestimation of active MSR cell abundance [START_REF] Colangelo-Lillis | Diversity decoupled from sulfur isotope fractionation in a sulfatereducing microbial community[END_REF] and/or underestimation of the rate of sulfide oxidation [START_REF] Jørgensen | Sulfur biogeochemical cycle of marine sediments[END_REF][START_REF] Treude | Rapid sulfur cycling in sediments from the Peruvian oxygen minimum zone featuring simultaneous sulfate reduction and sulfide oxidation[END_REF] might lead to a large underestimation of csSRR. It is possible that the current observed range of csSRR in natural environments is biased towards low rates [START_REF] Jørgensen | Sulfur biogeochemical cycle of marine sediments[END_REF], whereas csSRR is directly measured in pure culture experiments that allow the entire range of possible MSR-induced fractionations to be explored (Sim et al., 2011a,b) (Fig. 5).

We consider that pure culture experiments linking fractionation factors ε MSR and λ MSR to csSRR provide a valid first order approximation of the behavior of sulfur isotope fractionation to MSR. We used a nonlinear regression model in combination with pure culture experiment ε MSR , λ MSR and csSRR data to model the full range of MSR fractionation potentially expressed in natural environments (see supplementary 4.2 for detailed explanation). This approach does not consider the other environmental controls on MSR induced isotopic fractionation and should not be used to directly deduce a csSRR from our data but provide a simplistic way to model the ε MSR and λ MSR co-variation produced by MSR.

Results and discussion

Iron sulfides are found in all investigated gypsum and marl samples from VdG (Fig. 1A, supplementary section 3.1, and previously by [START_REF] Panieri | Microbial communities in Messinian evaporite deposits of the Vena del Gesso (northern Apennines, Italy)[END_REF]Sinninghe Damsté et al. (1995)). Iron sulfides form from the reaction of hydrogen sulfide produced by MSR with reactive iron, thus providing evidence for an active biogeochemical sulfur cycle during the first phase of the MSC. Data points represent the 33 S vs δ 34 S of sulfate and sulfide of gypsum (massive, banded, and branching selenite) and marl samples from the PLG unit of the VdG Basin. The dashed black arrow points to the isotopic composition of the first sulfide formed from oceanic sulfate by MSR considering the most depleted sulfide of marls measured in this study as a representative endmember. The curved arrow represents the calculated trajectory between this first sulfide and a pooled sulfide formed after the complete consumption of the initial sulfate reservoir. The blue line represents mixing of Messinian oceanic sulfate (SW) with the average isotopic composition of modern rivers (Freshwater), with f riv = 0r e p r e s e n t i n g the marine endmember and f riv = 1 representing the riverine endmember. Note that the contribution of riverine SO 4 to gypsum SO 4 isotopic composition seems negligible. (For interpretation of the colors in the figure, the reader is referred to the web version of this article.)

Marl interbeds

In marls, iron sulfides display the smallest variability, the lowest δ 34 S IS (average -32 ±5 vs CDT (1σ )) and the highest 33 S IS (average 0.12 ± 0.04 (1σ )) of all analyzed samples (Fig. 1A, Fig. 4, supplementary section 3.1). These isotopic values are typical of environments where the amount of available dissolved sulfate for MSR is effectively infinite (e.g. [START_REF] Gomes | Sulfur isotope fractionation in modern euxinic systems: implications for paleoenvironmental reconstructions of paired sulfatesulfide isotope records[END_REF][START_REF] Johnston | Sulphur isotopes and the search for life: strategies for identifying sulphur metabolisms in the rock record and beyond[END_REF]. We calculated net isotope fractionation factors 34 ε net and 33 λ net for marl samples considering a sulfate isotopic composition equal to that of the Messinian Ocean, based on the assumption that sulfide in marls was produced via the reduction of oceanic sulfate.

Marls fractionation factors fall within the MSR field (Fig. 5) and in fact represent the set of samples in this study that are the closest to the maximal fractionation expressed during MSR, with similar values to what is observed in modern organic-rich hypoxic to euxinic basins (e.g. [START_REF] Gomes | Sulfur isotope fractionation in modern euxinic systems: implications for paleoenvironmental reconstructions of paired sulfatesulfide isotope records[END_REF][START_REF] Johnston | Sulphur isotopes and the search for life: strategies for identifying sulphur metabolisms in the rock record and beyond[END_REF]. This is consistent with previous observations suggesting that these organic-rich marls formed under stratified and frequently euxinic water columns (Dela Pierre et al., 2014;Isaji et al., 2019;Lugli et al., 2010;Sinninghe Damsté et al., 1995) (Fig. 7a).

Gypsum beds

Petrographic evidence: co-occurrence of MSR and sulfide oxidation

Petrographic analyses of gypsum samples show that iron sulfides are associated with either marine snow floccules (Fig. 2J-L) or filamentous fossils (Fig. 2I-G), which are dominant components of the VdG gypsum (Natalicchio et al., 2022;[START_REF] Panieri | Microbial communities in Messinian evaporite deposits of the Vena del Gesso (northern Apennines, Italy)[END_REF]Schopf et al., 2012) and PLG deposits of other Mediterranean marginal basins (Dela Pierre et al., 2015;Natalicchio et al., 2022). The origin of the filaments has been disputed. They have been interpreted as cyanobacteria [START_REF] Panieri | Microbial communities in Messinian evaporite deposits of the Vena del Gesso (northern Apennines, Italy)[END_REF], even if they are generally larger (∼70 µm diameter) than most modern cyanobacterial filaments (Schopf et al., 2012 and references therein) (Fig. 2E-I). Moreover, the fact that filaments are mostly located both in the reentrant angle and along the vertical growth bands (Fig. 2D-F) suggests that they formed a benthic mat located at the gypsum-bottom water interface and that filamentous microorganisms were entrapped within gypsum in their living position, rather than after having settled from the photic zone to the basin floor (Dela Pierre et al., 2015;Schopf et al., 2012) (Fig. 7C). Finally, the presence of intracellular iron sulfide globules (Fig. 2G-I) -interpreted as resulting from diagenetic transformation of sulfur globules (Dela Pierre et al., 2015;Schopf et al., 2012) -is a diagnostic feature of benthic giant CSB. Such observations confirm the affiliation of the observed filament networks to chemotrophic microbial mats where both MSR and sulfide oxidation were active (Dela Pierre et al., 2015;Natalicchio et al., 2022;Schopf et al., 2012).

Isotopic composition of PLG gypsum deposits 5.2.2.1. Influence of syn-depositional or early diagenetic processes on multiple sulfur isotopes

The sulfur isotopic composition of the dissolved sulfate ion in equilibrium with the VdG gypsum ranged between 20.73 and 24.11 for δ 34 S SO4 , and between -0.001 and 0.049 for 33 S SO4 (Fig. 1A, 3, 4 Supplementary Table S2). These values fall within a narrow range that is centered around the isotopic composition of the Messinian Ocean dissolved sulfate ion (δ 34 S MESS = 22.2 ± 0.2 and 33 S MESS = 0.035 ± 0.013 ) (Masterson et al., 2016) (Fig. 1A, 3, 4). In contrast to the almost invariant isotopic composition of the sulfate, the iron sulfides entrapped in the same gypsum samples display a large range of isotopic compositions (-36 < δ 34 S IS < +9 and -0.017 < 33 S IS < 0.149), from high 33 S ISlow δ 34 S IS , which is comparable to that of iron sulfides contained within marl deposits intercalated to the gypsum beds, to close to zero 33 S IS and δ 34 S IS (Fig. 1A, Fig. 4, Supplementary section3.1).

The theoretical maximal sulfur isotopic fractionation induced by MSR (ε MSR ≈-70 ) generates sulfide isotope values of ∼-50 in δ 34 S IS when the sulfate used during MSR has a marine isotopic composition (∼+20 ). In some cases, the isotopic composition of iron sulfide differs substantially from this maximal estimate due to sulfide pooling controlled by local diagenetic or sedimentological processes (e.g. [START_REF] Masterson | Interpreting multiple sulfur isotope signals in modern anoxic sediments using a full diagenetic model (California-Mexico margin: Alfonso Basin)[END_REF]Pasquier et al., 2021bPasquier et al., , 2017)).

Sulfide pooling occurs when the rate at which sulfate is consumed (and sulfide is produced) in a given environment overcomes the rate at which the sulfate pool is replenished by diffusion from the overlying sulfate-rich water column [START_REF] Fike | Rethinking the ancient sulfur cycle[END_REF]. This leads to the progressive enrichment in 34 S of the residual sulfate and produces sulfide pools following a Rayleigh-type distillation. The observed isotopic difference (ε net ) between oceanic sulfate values and sedimentary sulfide is therefore smaller than the actual MSR induced fractionation (ε MSR ) because the pool of sulfate used by MSR is not oceanic sulfate (and therefore replenishable) but an aliquot of it (and therefore finite). For example, sulfide pooling has been proven to control the isotopic composition of sulfides in cases where sedimentation rates are so high that porewater sulfate is disconnected from the overlying water column (Pasquier et al., 2017), or where high concentrations or reactivities of organic matter in sediments results in efficient sulfate depletion due to MSR (Pasquier et al., 2021b). In other cases, iron sulfides enriched in 34 S are observed due to sediment mixing by oceanic currents or bioturbation bringing together pooled 34 S-enriched sulfide with newly formed 34 S-depleted sulfide [START_REF] Fike | Rethinking the ancient sulfur cycle[END_REF]Pasquier et al., 2021a).

The petrographic evidence and isotopic signature of gypsum and co-occurring iron sulfides from the VdG deposits infers particular depositional settings that exclude the role of sulfide pooling in shaping the sulfur isotopic signatures measured in this study. We detail the reasoning in the following paragraphs.

In the δ 34 S vs 33 S space, part of the gypsum iron sulfide samples fall along a curve (see the curved arrow in Fig. 4) consistent with the mixing of two sulfur reservoirs (Ono et al., 2006): (i) sulfide formed by MSR from an infinite, undistilled sulfate reservoir where the isotopic fractionation between sulfate and sulfide is close to equilibrium (corresponding here to the most depleted sulfide measured in this study with δ 34 S ∼-40 and positive 33 S ∼+0.17 ) and (ii) what would seem to be pooled sulfide produced by the consumption of SO 4 by MSR (and therefore characterized by an isotope composition that tends to approach that of initial, unreacted sulfate, as sulfate is consumed) (Ono et al., 2006;[START_REF] Zhang | Redox chemistry changes in the Panthalassic Ocean linked to the end-Permian mass extinction and delayed Early Triassic biotic recovery[END_REF].

To test the possible influence of sulfide pooling on the isotopic composition of our samples, we applied a Rayleigh type distillation to estimate the relative amount of sulfate necessary to be consumed to reproduce our data. We calculated that 55 to 95% of the sulfate reservoir should be consumed by MSR to reproduce the 34 S-enrichment in our samples in which δ 34 S IS falls between -20 and +10 (capturing 50% of the iron sulfides measured) (Fig. 4). We tested this with fractionation factors of -70 (i.e., the thermodynamic equilibrium fractionation between sulfate and hydrogen sulfide) and -58 (i.e., the net fractionation between the lowest measured δ 34 S IS and its corresponding δ 34 S SO4 ).

Such significant consumption of sulfate should either: (i) have led to the undersaturation of bottom waters with respect to gypsum, thus resulting in partial dissolution of the underlying gypsum, or (ii) have produced a substantial 34 S-enrichment of the isotopic signature of the residual sulfate pool preserved in gypsum. However, detailed petrographic studies of VdG gypsum here and in both Reghizzi et al. (2018) and [START_REF] Panieri | Microbial communities in Messinian evaporite deposits of the Vena del Gesso (northern Apennines, Italy)[END_REF] indicate that gypsum laminae have grown continuously without being interrupted by dissolution episodes (Fig. 2) and no substantial 34 Senrichment of the gypsum phase (Figs. 1, 3, and 4) have been observed in our samples, consistently with Lugli et al. (2007).

Finally, a drastic drop in the sulfate concentration within CSB mats seems highly unlikely because (i) mats hosting CSB grew on gypsum lamina (Fig. 2) that would provide an effectively infinite source of sulfate upon gypsum undersaturation; however, no interlaminae dissolution features are observed as stated above, and (ii) CSB can efficiently oxidize reduced sulfur to SO 4 [START_REF] Brüchert | Biogeochemical and Physical Control on Shelf Anoxia and Water Column Hydrogen Sulphide in the Benguela Coastal Upwelling System off Namibia[END_REF]Callbeck et al., 2021;[START_REF] Treude | Rapid sulfur cycling in sediments from the Peruvian oxygen minimum zone featuring simultaneous sulfate reduction and sulfide oxidation[END_REF].

The combined action of biogeochemical production of SO 4 by CSB, together with an infinite SO 4 reservoir (represented by the gypsum substrate below the mat and the water column above the mat), exclude the possibility of a large reservoir effect due to sulfate consumption. We therefore conclude that pooling effects did not modify the isotopic composition of our samples.

Sulfide and sulfate mineral formation from the same sulfur pool

An alternative scenario is therefore needed to explain the large range of variation in δ 34 S isotopic composition of iron sulfides entrapped in gypsum (-58 < ε net < -14 ). Previous nano-SIMS studies have shown that the isotopic composition of sulfides measured within microbial mats varied with depth due to variable sulfate concentration and/or csSRR within the mat (e.g. [START_REF] Fike | Micron-scale mapping of sulfur cycling across the oxycline of a cyanobacterial mat: a paired nanoSIMS and CARD-FISH approach[END_REF]. The isotopic compositions reported here, however, represent bulk measurements for each PLG gypsum cycle, and therefore are reflective of the most prominent processes governing the partitioning of sulfur species in the mat within several "generations" of CSB hosting mat (rather than local intra-mat conditions).

In the PLG deposits studied here, gypsum is a bottom-grown precipitate and is composed of an alternation of clear and turbid laminae (Fig. 2D-F). The mm-thick turbid laminae entrap filaments of CSB in close association with marine snow flocules rich in organic matter and iron sulfide minerals, (Fig. 2D-L). This suggests that sulfate reduction and sulfide re-oxidation by CSB bacteria occurred in the same space, most probably in the mat adhering to the previously formed gypsum laminae (Fig. 2D-F, Fig. 7B,C). Since the surface underlying each successive layer of microbial mat was gypsum, the bottom boundary of the MSR zone was restricted to the base of the microbial mat, and sulfides were produced insitu with no contribution of hydrogen sulfide diffusing from below (Fig. 7C).

Accordingly, iron sulfides co-precipitated with gypsum have multiple sulfur isotopic values ( 33 S IS -δ 34 S IS ) that neither reflect the overall water column conditions of the basin nor the bottom water sulfur cycling, but instead provide information on the dynamic of sulfur cycling within the mats permineralized in gypsum.

We interpret the 33 S IS -δ 34 S IS signature of iron sulfides as reflecting isotope fractionation processes associated with the biogeochemical sulfur cycle during gypsum precipitation (Fig. 7C). The same pool of sulfate was used to precipitate gypsum and to produce hydrogen sulfide through MSR. It is important to note that the particular depositional conditions under which the VdG iron sulfide and gypsum minerals co-precipitated are drastically different from what is observed in traditional marine sediments settings (Fig. 7B,C).

Paired sulfate-sulfide isotopes as a fingerprint of reoxidative sulfur cycling

The combination of iron sulfide and gypsum co-precipitation and the absence of sulfide pooling influence here allows for the calculation of net fractionation factors ( 34 ε net -33 λ net ) between sulfate from gypsum and their coeval iron sulfides in our samples. These net sulfate-sulfide isotope fractionation factors fall between the range of isotopic fractionations observed in pure cultures of microbial sulfate reducers and those observed in pure cultures of sulfur disproportionators (Fig. 5). Therefore, we conclude that a process characterized by 33 λ higher than maximal 33 λ MSR (>0.515) contributed, together with MSR, to define the isotopic composition of at least a part of the gypsum samples. The candidate processes include biotic H 2 S oxidation ( 33 λ H2S-SO4ox = 0.513-0.529) and intermediate oxidation state sulfur disproportionation ( 33 λ H2S-SO4disp = 0.514-0.528) (see Supplementary Table S3).

In support of this idea, the ten most 34 S-enriched samples (δ 34 S IS 8 between -20 and +8 and 33 S IS between 0.01 and 0.05 ) deviate from the mixing curve (Fig. 4) that assumes sulfide originating from a reservoir affected by sulfide pooling. Indeed, the contribution from a pooled sulfide reservoir is highly unlikely (see section 5.2.2.1), implying that this deviation is due to an additional sulfide pool. This alternative sulfide pool has a higher 33 S IS than seawater (0.035 ), and was possibly produced by disproportionation, as proposed in studies where a similar deviation was observed [START_REF] Zhang | Redox chemistry changes in the Panthalassic Ocean linked to the end-Permian mass extinction and delayed Early Triassic biotic recovery[END_REF].

This scenario is fully supported by the occurrence of two distinct sulfide pools in gypsum (Fig. 2, Fig. 4): the first one, associated with marine snow floccules (Fig. 2J-L), was produced by MSR in an infinite sulfate reservoir; the second, associated with remains of sulfide oxidizing bacteria (Fig. 2G-I), results from the re-oxidative part of the sulfur cycle.

In support of the significant occurrence of re-oxidative sulfur cycling, the sulfur and oxygen isotopic composition of the dissolved sulfate ion in equilibrium with the VdG gypsum (20.73 < δ 34 S SO4 vs CDT < 24.11; -0.001 < 33 S SO4 vs CDT <0.049 ; 8.01 <δ 18 O SO4 vs SMOW < 10.84) (Fig. 1, 3, 4 Supplementary Table S2) fall within a restricted range centered around the isotopic compositions of the dissolved sulfate ion of the Messinian Ocean (δ 34 S MESS = 22.2 ± 0.2 and 33 S MESS = 0.035 ± 0.013 ) (Masterson et al., 2016) and δ 18 O MESS = 9 ± 2 (Markovic et al., 2016). This implies negligible contributions from freshwater sulfate (Fig. 1,[START_REF] Lein | Isotopic composition of sulfur compounds in sediments of the Pacific[END_REF]4). Together with the very low iron sulfide concentration in gypsum (Supplementary section 3.1), the lack of 34 S and 18 O enrichment of sulfate characteristic of MSR [START_REF] Kah | Geochemistry of a 1.2 Ga carbonateevaporite succession, northern Baffin and Bylot Islands: implications for Mesoproterozoic marine evolution[END_REF][START_REF] Melegy | Microstructure and geochemistry studies on Messinian gypsum deposits from the Northern Coast of Egypt[END_REF] implies either that MSR activity during gypsum deposition was limited, or that the re-oxidation of sulfide into sulfate was nearly complete, i.e. the sulfur cycle was cryptic (Aller et al., 2010;Blonder et al., 2017) (Fig. 3).

Modeling multiple sulfur isotopes in the Vena del Gesso basin

Modeling strategy

To evaluate the importance of microbially-induced sulfide oxidation and to test the hypothesis of cryptic sulfur cycling, we quantified the relative contribution of biogeochemical processes in defining the multiple S isotope compositions of the VdG samples applying the numerical model described in section 4.3. In order to find a best fit to our data we tested several scenarios of sulfur redox cycling (see supplementary section 9), each one characterized by different relative sulfur fluxes (controlled by dimensionless f x parameters defined in supplementary section 5) and by a combination of distinct isotope fractionations (see Table S3) according to the following scenarios: cryptic sulfur cycling. The average rate of vertical gypsum accumulation at VdG was 1.4 mm.yr -1 (Reghizzi et al., 2018). Given that our model results imply that more than 80% of sulfate constituting this gypsum was first reduced by MSR, we calculated that approximately 1.9 mmol.cm -2 .yr -1 of sulfate was consumed by MSR. Then, by considering a rough C:S stoichiometry of 2:1 for the MSR reaction [START_REF] Jørgensen | Sulfur biogeochemical cycle of marine sediments[END_REF], we estimated a gross benthic organic carbon flux of 3.8 mmol.cm -2 .yr -1 remineralized by MSR during PLG gypsum formation. This is consistent with the large amounts of organic matter contained in the VdG interbeds. Marl deposits have a total organic carbon (TOC) content comparable to that of euxinic environments [START_REF] Jovovic | Quantitative and specific recovery of natural organic and mineral sulfur for (multi-)isotope analysis[END_REF]Lugli et al., 2007;Sinninghe Damsté et al., 1995), whilst TOC concentrations in gypsum range between 0.1 to 0.96% (Lugli et al., 2007), which is substantial considering the enormous dilution effect arising from gypsum depositional rates that are 25 to 250 times greater than marls (Reghizzi et al., 2018)). The benthic carbon flux involved in MSR we estimate for the VdG basin is comparable to that occurring in the world's most productive ocean regions, such as the Benguela and Peru-Chile margins, where the organic carbon remineralized through MSR ranges between 0.2 and 20 mmol.cm -2 .yrs -1 [START_REF] Bowles | Global rates of marine sulfate reduction and implications for sub-sea-floor metabolic activities[END_REF][START_REF] Treude | Rapid sulfur cycling in sediments from the Peruvian oxygen minimum zone featuring simultaneous sulfate reduction and sulfide oxidation[END_REF]. In these areas, dense mats populated by microbial sulfate reducers and CSB, similar to the fossil microbial assemblages found in the VdG gypsum, fuel a cryptic sulfur cycle [START_REF] Brüchert | Biogeochemical and Physical Control on Shelf Anoxia and Water Column Hydrogen Sulphide in the Benguela Coastal Upwelling System off Namibia[END_REF]Callbeck et al., 2021;[START_REF] Treude | Rapid sulfur cycling in sediments from the Peruvian oxygen minimum zone featuring simultaneous sulfate reduction and sulfide oxidation[END_REF].

The above considerations do not necessarily imply that the VdG basin was characterized by similar primary productivity to modern oxygen minimum zones but, rather, by comparable benthic carbon fluxes and carbon and sulfur cycling rates. Indeed, aerobic consumption of organic carbon in the VdG basin may have been limited to the upper layer of the water column due to density stratification (Natalicchio et al., 2022), leading to high organic carbon export fluxes (Pellegrino et al., 2021) and large quantities of benthic organic carbon available for MSR (Fig. 7). Conversely, very high primary productivity controls the export flux of organic carbon in modern productive margins.

Our findings support a recent model of PLG deposition [START_REF] Aloisi | The geochemical riddle of "low-salinity gypsum" deposits[END_REF] which suggests that the activity of benthic CSB mats drove increases in dissolved sulfate concentration that triggered gypsum precipitation. The inhibition of gypsum precipitation in deeper, more distal lateral equivalents of PLG gypsum (Natalicchio et al., 2019;[START_REF] Sabino | The response of water column and sedimentary environments to the advent of the Messinian salinity crisis: insights from an onshore deep-water section (Govone, NW Italy)[END_REF], together with our results showing that more than 80% of the sulfate constituting the more marginal VdG gypsum has been cycled through a dynamic sulfur cycle, suggest that gypsum precipitation could have been controlled by the efficiency of the re-oxidative portion of the sulfur cycle.

Concluding remarks

Our work suggests a tight coupling of microbially driven sulfate reduction and sulfide oxidation in the Vena del Gesso basin.

Primary Lower Gypsum deposits accumulated during the first phase of the Messinian Salinity Crisis and characterized by identical lithofacies, stacking patterns, petrographic features and microbial fossils are widespread across the Mediterranean (see Fig. 1C) (Lugli et al., 2010;Natalicchio et al., 2022;Pellegrino et al., 2021). Whereas these deposits were previously thought to be restricted to semi-enclosed basins with a water column depth of no more than 200 m (Lugli et al., 2010), recent studies have demonstrated that some were deposited at depths of up to 800 m and in basins not necessarily restricted by a sill (Ochoa et al., 2015;Raad et al., 2021).

If the cryptic sulfur cycle we describe was also active during the deposition of gypsum elsewhere in the Mediterranean, this would imply that the restriction of the Atlantic connection at the end of the Miocene led to not only the widespread deposition of evaporites, but also significantly altered the short-term carbon and sulfur cycles of the Mediterranean [START_REF] Waldeck | Calibrating the triple oxygen isotope composition of evaporite minerals as a proxy for marine sulfate[END_REF]. Further, the extensive occurrence of microbial mats populated by CSB on modern high-productivity continental margins or restricted basins, together with the results of the present study, suggest that cryptic sulfur cycling might have been prevalent in analogue environments of the geological past.

CRediT authorship contribution statement

The study was designed by Giovanni Aloisi and Laetitia Guibourdenche. Samples were collected by Laetitia Guibourdenche, Giovanni Aloisi, Marcello Natalicchio and Francesco Dela Pierre. Multiple sulfur isotope acquisition was performed by Laetitia Guibourdenche under the supervision of Pierre Cartigny. Photomicrographs were taken by Marcello Natalicchio. Laetitia Guibourdenche and Giovanni Aloisi designed the numerical model and interpreted the data. Laetitia Guibourdenche led the writing of the manuscript with contributions from all coauthors.

Figure 1 . 1

 11 Figure 1. 1 Schematic representation of carbon remineralization processes in the various biogeochemical zones within marine sediments that are characterized by electron acceptors. FromJørgensen and Kasten (2006), modified fromFroelich et al. (1979). Note that the methane (CH4) and sulfate (SO4 2-) concentration curves cross each other at the lower limit of the sulfate reduction zone(Jørgensen and Kasten, 2006).

Figure 1 . 2

 12 Figure 1. 2 Schematic representation of the effect of the weathering of three key types of rocks, where arrows represent the stoichiometry of the reaction. Carbonate weathering (green) releases calcium and alkalinity from the rock reservoir and CO2 from the atmosphere to the ocean. Carbonate deposition (black) reverses this process and decreases alkalinity and calcium concentration and releases CO2 to the atmosphere. Silicate weathering (purple) transfers calcium from the rock reservoir to the ocean and two moles of carbon from the atmosphere to the ocean. Evaporite weathering (orange) transfers only calcium to the ocean. FromShields and Mills (2020).

Figure 1 . 3

 13 Figure 1. 3 Scheme representing the main sulfur redox processes, i.e MSR (green arrows) and sulfide reoxidation (purple arrows) and their global annual fluxes in marine sediments.Fossing and Jørgensen (1990);[START_REF] Jørgensen | A Thiosulfate Shunt in the Sulfur Cycle of Marine Sediments[END_REF];[START_REF] Thamdrup | Thiosulfate and Sulfite Distributions in Porewater of Marine-Sediments Related to Manganese, Iron, and Sulfur Geochemistry[END_REF].These redox processes occur mostly in small restricted marine reservoirs, generally in the porewater environment, but can also occur within the water column in specific cases. In any case, these restricted reservoirs have limited exchanges with the global ocean (blue arrow). The fluxes of sedimentary sulfide and sulfate minerals being buried in the rock record (brown arrows) are the net result of the redox processes. While the orders of magnitude for the fluxes shown are well known, their precise estimation depends on the method used to calculate, as well as on the time period for which the flux is integrated and if the oceanic sulfur cycle is considered at steady state. a corresponds to the estimated flux of sulfate burial, considering that no major evaporitic basins are actually precipitating sulfate bearing salts and that barite burial (∼0.02 Tmol.years-1) is negligible compared to the fluxes considered here(Carter et al., 2020; Dickens et al., 2003) b Estimated fluxes considering sulfur cycle at steady state and integrating the flux of SO4 buried into sulfate bearing minerals over the course of the last 20Myrs(Tostevin et al., 2014) C Estimated flux after(Berner, 1982) d,e,f,g Range of existing estimates of annual fluxes of sulfate being reduced by MSR via various methods[START_REF] Bowles | Global rates of marine sulfate reduction and implications for sub-sea-floor metabolic activities[END_REF][START_REF] Bradley | Widespread energy limitation to life in global subseafloor sediments[END_REF] Canfield et al., 2005; Thullner et al., 2009). h,i,j Reoxidative fluxes of hydrogen sulfide to sulfate calculated using estimates from d,e,f and followingFossing and Jørgensen (1990);[START_REF] Jørgensen | A Thiosulfate Shunt in the Sulfur Cycle of Marine Sediments[END_REF];[START_REF] Thamdrup | Thiosulfate and Sulfite Distributions in Porewater of Marine-Sediments Related to Manganese, Iron, and Sulfur Geochemistry[END_REF], that consider that 90% of the sulfate being processes by MSR is ultimately reoxidized back to sulfate k Reoxidative fluxes of hydrogen sulfide to sulfate considering that numbers given in references b,c,d,e,f represent valid estimations and steady state i.e., the flux of sulfur that is not buried as sedimentary sulfide is necessarily reoxidized. In that case, the 90% of sulfide reoxidation estimation, widely used in the literature (see h,i, jfor references) is not valid anymore for the upper range of sulfate reduced by MSR estimation and up to 99% of the sulfide produced by MSR should have been reoxidized to satisfy mass balance.

Figure 1 . 4

 14 Figure 1. 4 Scheme representing the link between short-term and long-term sulfur cycles. The fluxes of sulfur exiting the restricted reservoir (sedimentary sulfide and sulfate burial) largely control long term change in oceanic sulfur cycle and are represented in brown. Changes in the proportion of the fluxes of sulfur being buried in the sediment (Tostevin et al., 2014) (brown arrows) and returned back to ocean through weathering (Burke et al., 2018) (dark blue arrows) control the global ocean sulfate concentration and impact the surface earth oxidation state. b Estimated fluxes considering sulfur cycle at steady state and integrating the flux of SO4 buried into sulfate

Figure 1 . 5

 15 Figure 1. 5 Geographic repartition of salt giants, modified by J. Lofi afterWarren (2010)

Figure 1 .

 1 Figure 1. 6 A. Age and approximate amount of CaSO4 presently sequestered as evaporites. B. Estimates of the amount of CaSO4 eroded since its initial deposition as evaporite. C. Estimate of the amount of CaSO4 originally precipitated. From Wortmann and Paytan (2012).

Chapter 1 -

 1 General Introduction _________________________________________________________________________________________________ 16

Figure 1 . 7

 17 Figure 1. 7 Mineral precipitation sequence in a concentrating brine, from Warren (2010) replotted from Briggs (1958).

  Figure 1. 8 Major elements constituting salinity in 1kg of sea water https://salinity.oceansciences.org/search.htm

Figure 1 .

 1 Figure 1. 10 Number of publications mentioning the term "Messinian salinity crisis" each year from 1976 to today in the Title, abstract or key words of published peer-reviewed papers (n=1167) from https://wwwscopus-com.

Chapter 1 -

 1 General Introduction _________________________________________________________________________________________________ 21

Figure 1 .

 1 Figure 1. 11 Map of the Mediterranean region showing two of the possible paleoenvironmental configurations proposed to have featured the Mediterranean during the MSC. The top panel represent the "deep desiccated basin" scenario fromHsü et al. (1973) and the bottom one represents "deep water, deep basins"[START_REF] Andreetto | Alternate interpretation of the Messinian salinity crisis, controversy resolved?[END_REF]. Modified fromKrijgsman et al. (2018)

  Figure 1. 12 Two (simplified) conceptual scenarii derived from the two step model of[START_REF] Andreetto | Alternate interpretation of the Messinian salinity crisis, controversy resolved?[END_REF] and accounting for thick MSC halite deposition in the deep basin. Stage 1(5.96-5.60 Ma, Krijgsman et al., 1999): primary gypsum accumulated in silled marginal basins. Stage 2: Scenario 1: moderate base level drop in the Mediterranean and evaporite emplacement by density currents in a stratified deep water environment[START_REF] Roveri | The Mediterranean Messinian salinity crisis: an Apennines foredeep perspective[END_REF] 2009; 2014b); Scenario 2: drying out of the Mediterranean basin highly concentrated in salts (brine pool) by evaporative shrinking[START_REF] Ryan | Modeling the magnitude and timing of evaporative drawdown during the Messinian salinity crisis[END_REF] Lofi et al., 2011b). Evaporite emplacement starts in deep water environment and ends in shallow water environment; During the MSC, continental margins and slopes underwent intense subaerial/subaqueous erosion. In red, the Messinian Erosion Surface (MES) where margins and slopes were subaerially exposed. Figure and captions modified from Umbrella proposal of the Deep-Sea Record of Mediterranean Messinian Events (DREAM) multi-phase drilling project.

1. 5

 5 Restriction of Atlantic-Mediterranean hydrological exchanges before the onset of the MSC During the Miocene (23.03 to 5.33 Ma), the closure of the Neo-Tethys ocean due to the drifting of African, Arabian and Indian plates restricted the main Tethyan water mass into two main bodies of water : the Mediterranean Sea and the Paratethys (the latter was later further divided into the modern Aral, Caspian, and Black seas) (Potter and Szatmari, 2009; Rögl, 1999)(Figure 1. 13). Since around 11Ma, the current region of Gibraltar has been the only point of hydrological connection of the Mediterranean Sea with the open ocean (Rögl, 1999) ( Figure 1. 13).

Figure 1 .

 1 Figure 1. 13 Paleogeographic reconstruction of the late Miocene configuration of the Mediterranean Sea and

Figure 1 .

 1 Figure 1. 15 Schematic representation of Mediterranean circulation before and after the start of the restriction of water masses exchanges with the Atlantic. From Bulian et al. (2022), adapted from(Alhammoud et al., 2010; Pinardi and Masetti, 2000;[START_REF] Topper | The precessional phase lag of Messinian gypsum deposition in Mediterranean marginal basins[END_REF]

Figure 1 .

 1 Figure 1. 16 Sketch showing the orbitally controlled paleoenvironmental changes affecting the marginal Piedmont basin (Northern Italy) before (A and B) and after (A and C) the onset of the MSC. Blue arrows indicate inputs of freshwater (thin arrows) and sea water from the open Mediterranean (thick arrows). Blue dashed arrows indicate water mixing and brown arrows aeolian transport. FromNatalicchio et al. (2019)

Chapter 1 -

 1 General Introduction _________________________________________________________________________________________________ 30

Iand 5 .

 5 co-authored during the course of my PhD .Appendix A corresponds to the tables of DSDP and IODP sulfur and oxygen isotopic data that I measured during my PhD and that are discussed in chapter 4 Appendix B corresponds to a compilation of δ 34 S isotopic values from both modern and ancient early sedimentary iron sulfides from the literature that was used in Chapter 4. Appendix C corresponds to a publication led by my PhD advisor Giovanni Aloisi in which I am second author: "The geochemical Riddle of low salinity gypsum deposits" published in Geochemica and Cosmochimica Acta in April 2022. In this paper we combined measurement of the water isotopic composition of the gypsum hydration water to measurements of fluid inclusion salinity and sulfur and oxygen isotope composition of gypsum. This was done in order to trace the water masses involved in the formation of gypsum during the MSC and propose a new model for the formation of these deposits that were formed under salinities lower than those expected if gypsum formed by evaporation only. I was mostly involved in measuring the isotopic composition of gypsum hydration Nature Reviews Earth & Environment | https://doi.org/10.1038/s43017-022-00273-0 | Published online xx xx xxxx NATURE REVIEWS | EARTH & ENVIRONMENT

Figure 1 .

 1 Figure 1. 19 Exaggerated diagram in the δ 34 S-δ 33 S space showing the curves associated to processes fractionating sulfur isotopes following slightly different mass dependent relationship (solid green and red lines) than the equilibrium fractionation line (solid blue). Their associated straight lines in the δ 34 S-∆ 33 S space corresponds to the dashed lines.

Figure 1 .

 1 Figure 1. 20 Exaggerated diagram showing the partitioning of sulfur isotopes during a mixing. The blue solid line represents the equilibrium fractionation line on which terrestrial sample should fall. The two crosses

Figure 1 .

 1 Figure 1. 21 Ranges of 34 ɛ and 33 λ values obtained in pure cultures of sulfur metabolizing microorganisms (shaded areas). Data from abiotic sulfide oxidation experiments comes from(Eldridge and Farquhar, 2018) (dark grey area with black dotted contour), data from pure cultures with sulfate reducers comes from(Bertran et al., 2018; Johnston et al., 2005; Leavitt et al., 2013; Sim et al., 2011a; and Sim et al., 2011b) (light grey area) and data from pure culture experiments with disproportionators comes from(Johnston et al., 2005) (white area with grey dashed contour). White squares represent the 34 ɛnet and 33 λnet measured in natural environments populated with biofilms dominated by chemolithotroph sulfide oxidizer(Magnuson et al., 2020; Zerkle et al., 2016).

Figure 2 . 1

 21 Figure 2. 1 Distribution of Messinian Evaporites in the Mediterranean. Map modified afterLofi (2018)

Chapter 2 -

 2 Material and Methods ___________________________________________________________________________________________ 46 erosion of MSC stage 1 gypsum, and Upper Gypsum (UG) deposits that were formed during the terminal or third phase of the MSC.

Figure 2 . 3

 23 Figure 2. 3 Correlation and astronomical tuning of PLG across the western Mediterranean basins. FromLugli et al. (2010)

Chapter 2 -

 2 Figure 2. 4 Compilation of the isotopic composition of gypsum from some western marginal basins of the Mediterranean (see legend). Figure fromGarcía-Veigas et al. (2018)

Chapter 2 -Figure 2 . 6 Figure 2 . 7

 22627 Figure 2. 6 Photomicrographs of filamentous fossils entrapped in PLG bottom grown selenite in (a) Nijar (Southern Spain) (b-e) Monte Tondo (Northern Apennines) (f) Crete (g) Cyprus. Transmitted light(a,b,c,d,f,g) and UV light (e). White triangles in (e) refers to non-fluorescent iron sulfides coating the filaments, of which the fluorescence suggest the presence of abundant organic matter. Figure fromNatalicchio et al. (2022).

Chapter 2 -Figure 2 . 8

 228 Figure 2. 8 Stratigraphic log of the VdG succession calibrated with the 100ka orbital eccentricity and precession curves[START_REF] Laskar | La2010: a new orbital solution for the long-term motion of the Earth[END_REF]

Figure 2 . 9

 29 Figure 2. 9 The 1 st and 2 nd PLG cycles in the Monticino quarry

Chapter 2 -Figure 2 .

 22 Figure 2. 10 Sampling group of the VdG PLG unit in the Monte Tondo quarry where cycles 3 to 16 have been sampled

  in each PLG cycle by a 4 to up to 30 m thick bed of primary gypsum (Figure 2. 8, Figure 2. 11 and Figure 2. 12). The gypsum beds of the first three cycles (PLG1 to PLG3) are composed exclusively of bottom-grown, vertically oriented, twinned selenite crystals (up to 2.5 m high)

Chapter 2 -Figure 2 .Figure 2 .

 222 Figure 2. 11 Picture showing the transition between PLG marls and gypsum that mark the beginning of cycle 4 (Monte Tondo section).

Chapter 2 -

 2 Figure 2. 13 Schematic representation showing the steps of growth of selenite that determine the typical orientation of the crystals and allow the entrapment of bacterial filaments in the re-entrant angle of the twins. Modified from(Lugli et al., 2010)

Chapter 2 -Figure 2 .

 22 Figure 2. 16 Banded selenite facies the small (cm to mm) crystals are separated by a thin carbonate layer

Figure 2 .

 2 Figure 2. 18 Schematic representation of the mode of formation of the branching selenite lithofacies from[START_REF] Natalicchio | Gypsum lithofacies and stratigraphic architecture of a Messinian marginal basin (Piedmont Basin, NW Italy)[END_REF]

2. 2 Figure 2 .

 22 Figure 2. 19 Distribution of Messinian Evaporites in the Mediterranean and location of sites sampled in this work (green star for Upper gypsum cores and pink star for the Vena del Gesso basin). Map modified afterLofi (2018)

Chapter 2 -Figure 2 .

 22 Figure 2. 20 Correlation and tuning of stage 3 deposits from marginal and intermediate depth basins.Figure fromAndreetto et al. (2021a)

  (Figure 2. 21). In the non-evaporitic interbeds of the same sections, shallow water brackish ostracods of Paratethyan affinity are found and display lower strontium isotopic values than the co-eval Messinian ocean (Figure 2. 21 and Figure 2. 22). This has been interpreted to reflect the dilution of the Mediterranean sub-basins with continental waters during marl deposition

Chapter 2 -Figure 2 .

 22 Figure 2. 21 Ideal UG cycle afterAndreetto et al. (2022a)

Chapter 2 -

 2 Material and Methods ___________________________________________________________________________________________ 69 tempting to draw a correlation between deposits having the same type of fossil contents (here between intermediate and marginal basins), the concomitant occurrence of extremely shallow water fauna in connected basins having several hundreds to thousands of meters of difference in depth is impossible. In Andreetto et al. (2022a) in the Caltanissetta basin of Sicily, we propose that the alternation of brackish shallow water fauna in marls and marine calcareous nannofossils assemblages in gypsum and their respective strontium isotopic variations, reflect hydrological oscillations of the basin between purely lacustrine settings and partially marine conditions during UG deposition (see Appendix E for full text). This scenario explains the occurrence of shallow water fauna in both intermediate and marginal basins by the temporal disconnection of their occurrence in each basin depth. In summary, the complexity of the paleontological record, the scarcity and often incomplete character of phase 3 deposits outcropping on land, have led to contrasting paleoenvironmental interpretations of the terminal phase of the MSC in the entire Mediterranean. The proposed scenarios range from an almost entirely refilled Mediterranean (after the massive precipitation of halite in the deep basins) fed by the Atlantic, with most of the basins connected to each other, to the occurrence of isolated endoreic basins fed only by freshwater inputs (rivers and/or Parathethys), or even an alternation between these two scenarios at the precessional scale (Figure 2. 23). In chapter 4 of this thesis we will propose an estimation of the water level at the time of UG deposition.

Figure 2 . 72 Figure 2 .

 2722 Figure 2. 23 Map of the Mediterranean region showing two of the possible paleoenvironmental configurations proposed to have featured the Mediterranean during stage 3. Modified fromKrijgsman et al. (2018)

Figure 2 .

 2 Figure 2. 26 Site 654. A. Core slab of a laminated gypsum consisting of fine grained graded granular gypsum with a clastic gypsum intercalation at the center. B. Thin section of a similar laminated rock. Modified from[START_REF] Lugli | The deep record of the Messinian salinity crisis: evidence of a non-desiccated Mediterranean Sea[END_REF]

Figure 2 .

 2 Figure 2. 27 Site 376. Core photograph of a gypsrudite. Rounded to angular pieces of clastic gypsum rest in a matrix of marlstone and fine grain clastic gypsum. Picture from[START_REF] Hsü | History of the Mediterranean salinity crisis[END_REF]

Figure 2 .Figure 2 .

 22 Figure 2. 28 Site 134. Laminated anhydrite with scattered nodules. Picture fromFriedman (2007)

Figure 2 .

 2 Figure 2. 32 Schematic cross section illustrating the evaporite facies distribution at DSDP sites 133 and 134. Figure fromRyan et al. (1973)

Figure 2 .

 2 Figure 2. 34 Map of the central Mediterranean area with location of site 374 A approximately at the intersection of the yellow lines within the yellow square on this map. From[START_REF] Camerlenghi | Uncovering the Mediterranean Salt Giant (MEDSALT)-Scientific Networking as Incubator of Cross-disciplinary Research in Earth Sciences[END_REF]

Chapter 2 -

 2 Material and Methods ___________________________________________________________________________________________ 82 Quaternary transition. This transition is highly irregular and describes V-shaped incisions. The samples analysed in this study correspond to sub-unit IIIb (core 16-20) and sub-unit IIIc (core 21-22). Sub-unit IIIb consists of cyclic alternation of barren dolomitic mudstones and gypsum layers (both clastic, anhydrite and primary cumulates; Lugli et al., 2015); sub-unit IIIc contains anhydrite and halite at the bottom, the latter of uncertain origin (similarly to halite cored in Site 134, see discussion above) (Figure 2. 31).

Figure 2 .

 2 Figure 2. 35 Interpretation of seismic profiles at sites 375 and 376 at the Florence Rise. Figure from Hsüet al. (1978c)

Figure 2 .

 2 Figure 2. 36 Scheme showing the principle of the sulfur extraction installation in IPGP to convert the sulfur contained in samples into silver sulfide. Modified fromMuller (2016)

  Figure 2. 37, following the well-established protocol, similarly to Labidi et al. (2012) or Au Yang et al. (2016).

Figure 2 .

 2 Figure 2. 37 Schematic representation of the conversion and purification line of SF6. The orange compartment corresponds to the final trap where excess fluorine is stored in the form of BrF5. Modified from original figure, courtesy of Arnaud Durverger.

Figure 2 .

 2 Figure 2. 39 Plot of the uncertainty associated with the determination of 33 λnet (σλnet) as a function of the range of fractionation between sulfate and sulfide. Here are plotted the calculated uncertainties for the Vena del Gesso dataset. Note the exponential relationship between σλnet and ε with σλnet increasing when the magnitude of the fractionation ε decreases.

Figure 2 .

 2 Figure 2. 40 Structure of the numerical model simulating the biogeochemical cycling of sulfur in an idealized marginal basin. The three boxes correspond to the inventories of sulfate (SO4 2-), sulfide (H2S) and sulfur with an intermediate oxidation state (sulfite: SO3 2 or elemental sulfur: S °). The arrows represent the following sulfur fluxes: Fin and Fout are the exchange fluxes with the ocean, Friv is the riverine flux, Fgyp is the gypsum precipitation flux, Fmsr is the flux of sulfur reduced by MSR, Fpy is the flux of sulfur being buried as pyrite, Fox is the flux of sulfur being reoxidised to sulfate or intermediate sulfur compound, Fdh is the flux of sulfur being disproportionated to sulfide, and Fdo is the flux of sulfur being disproportionated sulfate. Fluxes associated to biogeochemical transformations or mineral formation have specific fractionation factors 34 α and 33 λ defined in Table 2. 1 Sulfur fractionation factors used in the numerical model

  )) on the results of a pure culture experiment with microbial sulfate reducer Desulfovibrio vulgaris Hildenborough (DvH) under constant sulfate concentration and variable concentration of the electron donor lactate:

Figure 2 .

 2 Figure 2. 41 Nonlinear regression model fits using the equations from Leavitt et al., 2013 to major and minor sulfur isotope fractionations measured in experiments by Sim et al 2011 a and b. In both panels A and B paired sulfate-sulfide major 34 Ɛmsr or minor 33 λmsr observed fractionation factors are plotted against the recast csSRR rate of MSR activity.

Figure 3 . 1

 31 Figure 3. 1 PLG in the Mediterranean and in the VdG basin. A. Sulfur δ 34 SSO4 and δ 34 SIS isotopic curve of the PLG succession of the VdG basin, calibrated with the 100ka orbital eccentricity and precession curves[START_REF] Laskar | La2010: a new orbital solution for the long-term motion of the Earth[END_REF]. Note the nearly invariant δ 34 SSO4 isotopic variations compared to δ 34 SIS. B.Panoramic view of the 220m thick VdG succession constituted by an alternation of thick gypsum beds and marl intervals (PLG unit). C. Distribution of Messinian evaporites in the Mediterranean and the location of the VdG basin (pink star) (see Chapter 2 for references).

Figure 3 . 2 3 -

 323 Figure 3. 2 Macro-and microscopic characteristics of VdG PLG gypsum deposits. A. The basal part of cycle PLG5, composed of a basal dark marly interval followed by a thin whitish carbonate layer, in turn overlain by a gypsum bed consisting of dm-sized vertically oriented selenite crystals (massive selenite

3. 3 .

 3 1 δ 34 S and δ 18 O of sulfate

Figure 3 . 3

 33 Figure 3. 3 Stable sulfate isotopic composition of VdG gypsum. δ 34 SSO4 and δ 18 OSO4 isotopic composition of the dissolved sulfate ion in equilibrium with the PLG unit from VdG basin (dark pink) and the PLG unit of other marginal basins (light pink) including values from the Spanish marginal basins of Sorbas[START_REF] Evans | Coupled measurements of delta O-18 and delta D of hydration water and salinity of fluid inclusions in gypsum from the Messinian Yesares Member, Sorbas Basin (SE Spain)[END_REF] García-Veigas et al., 2018), Bajo Segura (García-Veigas et al., 2018), Murcia(Pierre, 1982) and Mallorca(García-Veigas et al., 2018), the Italian marginal Piedmont and Caltanissetta basins(El Kilany, 2018; García-Veigas et al., 2018) and the Polemi and Kalavasos/Psematismenos basins of Cyprus(Pierre, 1982)) (see section 2.3.3 for calibration of δ 34 SSO4 and δ 18 OSO4 isotopic data from this study and from the literature). The range of isotope compositions of sulfate-bearing minerals(Boschetti et al., 2011 and references therein;[START_REF] Claypool | The age curves of sulfur and oxygen isotopes in marine sulfate and their mutual interpretation[END_REF] Crockford et al., 2019; Ortí et al., 2010; Sakai, 1971; Utrilla et al., 1992) and of sedimentary sulfides(Balci et al., 2007; Burke et al., 2018) are represented as grey and light brown boxes. The possible range of isotopic composition of riverine inputs is bracketed by the isotopic composition of sulfate-bearing minerals and sedimentary sulfides representing the riverine endmember. Note that the contribution of riverine SO4 to gypsum SO4 isotopic composition seems negligible. The possible evolution of the isotopic composition of residual sulfate during MSR is indicated by the orange shaded area[START_REF] Fotherby | Modelling the Effects of Non-Steady State Transport Dynamics on the Sulfur and Oxygen Isotope Composition of Sulfate in Sedimentary Pore Fluids[END_REF]. The isotope composition of sulfate dissolved in top few centimeters of modern marine environments characterized by near complete reoxidation of H2S is represented by the blue box(Aller et al., 2010).

3. 5

 5 .2.1 Petrographic evidence: co-occurrence of MSR and sulfide oxidation Petrographic analyses of gypsum samples show that iron sulfides are associated with either marine snow floccules (Figure 3. 2J-L) or filamentous fossils (Figure 3. 2I-G), which are dominant components of the VdG gypsum

Chapter 3 -Figure 3 . 4

 334 Figure 3. 4 Multiple Sulfur isotopic composition of sulfate and sulfide minerals constituting PLG deposits. Data points represent the Δ 33 S vs δ 34 S of sulfate and sulfide of gypsum (massive, banded, and branching selenite) and marl samples from the PLG unit of the VdG Basin. The dashed black arrow points to the isotopic composition of the first sulfide formed from oceanic sulfate by MSR considering the most depleted sulfide of marls measured in this study as a representative endmember. The curved arrow represents the calculated trajectory between this first sulfide and a pooled sulfide formed after the complete consumption of the initial sulfate reservoir. The blue line represents mixing of Messinian oceanic sulfate (SW) with the average isotopic composition of modern rivers (Freshwater), with friv =0 representing the marine endmember and friv=1

Chapter 3 -

 3 Cryptic sulfur cycling during the formation of giant gypsum deposits _____________________________________________________________________________________________________ 134 3. 8). The same pool of sulfate was used to precipitate gypsum and to produce hydrogen sulfide through MSR. It is important to note that the particular depositional conditions under which the VdG iron sulfide and gypsum minerals co-precipitated are drastically different from what is observed in traditional marine sediments settings (Figure 3. 8B, C).

Figure 3 . 5

 35 Figure 3. 5 Comparison between fractionation factors determined in laboratory experiments and net fractionations deduced from VdG sulfide and sulfate minerals. Net fractionation factors 34 ɛnet and 33 λnet calculated for the gypsum (massive, banded, and branching selenite) and marl samples using the isotope compositions of the sulfate ion in equilibrium with gypsum and the sulfur isotope composition of iron sulfides (dots and crosses) compared with ranges of 34 ɛ and 33 λ values obtained in pure cultures of sulfur metabolizing microorganisms (shaded areas). For marls samples, where no gypsum precipitates occur, an isotope composition of the sulfate ion equal to the Messinian ocean value(Masterson et al., 2016) was set, based on the assumption that sulfides from marls were produced through the reduction of oceanic sulfate. See figure1.20 for references on data from abiotic sulfide oxidation experiments (dark grey area with black dotted contour), data from pure cultures with sulfate reducers (light grey area), data from pure culture experiments with sulfur disproportionators and 34 ɛnet and 33 λnet fractionation factors measured in natural environments populated with biofilms dominated by chemolithotrophic sulfide oxidizers (white squares).

Chapter 3 -

 3 Cryptic sulfur cycling during the formation of giant gypsum deposits _____________________________________________________________________________________________________ 136

Figure 3 . 6

 36 Figure 3. 6 Structure of the numerical model simulating the biogeochemical cycling of sulfur in an idealized marginal basin. The three boxes correspond to the inventories of sulfate (SO4 2-), sulfide (H2S) and sulfur with an intermediate oxidation state (sulfite: SO3 2 or elemental sulfur: S °). The arrows represent the sulfur fluxes described in section 3.4.3.1.

Chapter 3 -

 3 Cryptic sulfur cycling during the formation of giant gypsum deposits _____________________________________________________________________________________________________ 141 sulfur and then almost completely reoxidized back to sulfate (>85%) resulting in low net production of H2S (Figure 3. 8B,C) and confirming the existence of a "cryptic" sulfur cycle during gypsum formation.

Figure 3 . 8

 38 Figure 3. 8 Schematic representation of the biogeochemical conditions that accompanied PLG deposition. (A) Deposition of organic rich marls during humid phases (precession minima). (B) Deposition of gypsum in the margins and marls in the depocenter during arid phases (precession maxima). The relative size of the arrows represents the approximate magnitude of fluxes. (C) Details of a gypsum twin.

Figure 3 . 9 Chapter 3 -

 393 Figure 3. 9 Concentration of sulfides within gypsum and marls interbeds of the Vena del Gesso succession as a function of their δ 34 SCRS isotopic composition

Figure 3 .

 3 Figure 3. 10 Sensitivity test for varying values of fex of 0.96 (green area) and0.85 (brown area). The main differences from our best fit results (where fex =0.9) concern the sulfate isotopic values, with the area covered being smaller when the choosen fex is larger, possibly requiring a larger fmsr to explain the whole range of values covered by our sulfate data (see supplementary section 6 for further explanation on this point).

Chapter 3 -Chapter 3 -Figure 3 .

 333 Figure 3. 11 Measured and modelled Δ 33 SSO4 vs δ 34 SSO4 best fit with different values of fmsr. Individual curves represent the modelled isotopic composition sulfate data from gypsum samples for 0<fox<1 and 0<csSRR<100.

Chapter 3 -

 3 Cryptic sulfur cycling during the formation of giant gypsum deposits _____________________________________________________________________________________________________ 157 sulfide being oxidated into sulfate or intermediate sulfur compound) and ƐMSR (the extent of isotopic fractionation during MSR) were explored.

Figure 3 . 3 -

 33 Figure 3. 12 Measured and modelled Δ 33 SSO4 and Δ 33 SIS vs δ 34 SSO4 and δ 34 SIS data in A and 34 ɛnet and 33 λnet in B with the model fluxes and fractionation factors chosen for a scenario where there is no reoxidation of H2S. Individual curves represent the modelled isotopic composition in A and net fractionations factor for paired sulfate-sulfide data from the same gypsum samples in B for different values of fmsr and ƐMSR. fmsr is the fraction of the net incoming sulfur flux being removed from the marginal basin through microbial sulfate reduction rather than by gypsum accumulation and csSRR the extent of 34 S isotopic

Chapter 3 -Figure 3 .

 33 Figure 3. 13 Measured and modelled Δ 33 SSO4 and Δ 33 SIS vs δ 34 SSO4 and δ 34 SIS data in A and 34 ɛnet and 33 λnet in B with the model fluxes and fractionation factors chosen for the phototrophic reoxidation scenario. Individual curves represent the modelled isotopic composition in A and net fractionations factor for paired sulfate-sulfide data from the same gypsum samples in B for different values of fmsr and csSRR. fmsr is the fraction of the net incoming sulfur flux being removed from the marginal basin through microbial sulfate reduction rather than by gypsum accumulation and csSRR the cellular rate of microbial sulfate reduction. Mathematical relationship between the fluxes of sulfur Fx and the relative magnitude of these fluxes fx are given in section 5.

Figure 3 .

 3 Figure 3. 14 Measured and modelled Δ 33 SSO4 and Δ 33 SIS vs δ 34 SSO4 and δ 34 SIS data in A and 34 ɛnet and 33 λnet in B with the model fluxes and fractionation factors chosen for the abiotic oxidation of H2S into SO3 2- and its subsequent disproportionation. Individual curves represent the modelled isotopic composition in A and net fractionations factor for paired sulfate-sulfide data from the same gypsum samples in B for different values of fmsr and csSRR. fmsr is the fraction of the net incoming sulfur flux being removed from the marginal basin through MSR rather than by gypsum accumulation and csSRR the cellular rate of microbial sulfate reduction. Mathematical relationship between the fluxes of sulfur Fx and the relative magnitude of these fluxes fx are given in section 5.

4. 1 . 1

 11 Water level and basins depth in the Mediterranean after the acme of the MSC: an unsolved controversy

Chapter 4 -

 4 Late Messinian Mediterranean base level fall ___________________________________________________________________________________________ 162 the acme of the MSC, a further increase in salinity allowed km-thick halite sequences to accumulate, mostly in deep basins (>1000m deep) of the Mediterranean (Lofi, 2018; Roveri et al., 2014a) (Figure 4. 1).

Figure 4 . 1

 41 Figure 4. 1 Distribution of Messinian Evaporites in the Mediterranean and location of sites sampled in this work (red dots). Green dots represent DSDP/ODP sites where MSC evaporites have been found. Map modified after Lofi (2018).

Chapter 4 -

 4 Late Messinian Mediterranean base level fall ___________________________________________________________________________________________ 163

Stage 3

 3 deposits are interpreted to have either: (i) accumulated in isolated endorheic basins Chapter 4-Late Messinian Mediterranean base level fall ___________________________________________________________________________________________ 164

Figure 4 Figure 4 . 2

 442 Figure 4. 2) in the range of hundreds of meters, the former scenario implies that a significant

Chapter 4 -Figure 4 . 3

 443 Figure 4. 3 Lithological log of the cores studied in this work

  The individual paleogeographic and paleobathymetric context of each sub-basins of the deep Mediterranean from which the studied cores have been retrieved is difficult to assess. Their distance from freshwater sources, degree of restriction from the open ocean (Figure 4. 1) and Chapter 4-Late Messinian Mediterranean base level fall ___________________________________________________________________________________________ 173 their bathymetry (section 2.2.2) are and possibly were very different from each other, making

4. 3 . 3

 33 Isostasy based paleo-bathymetric estimation The paleobathymetry of stage 3 sedimentary layers initially deposited at depth Y' can be estimated by taking Y, the current depth of the top of the stage 3 deposits (i.e the Miocene-Pliocene boundary) and calculating the magnitude of various components of vertical motion (YY') (Figure 4. 2 and 4.4). The paleo-bathymetry (Y') can be calculated either (i) for a basin filled to global sea level (where the water column thickness is X'Y'full), compensating for isostasy due to Plio-Quaternary sediment load, sediment compaction and thermal subsidence (Scheme B of Figure Chapter 4-Late Messinian Mediterranean base level fall ___________________________________________________________________________________________ 176 4.4), or (ii) for a basin that has undergone significant water level drop (Scheme C of Figure 4.4), in which case isostatic rebound due to the removal of a significant part of the water column (XX') must also be taken into account (panel D of Figure 4.4). Indeed, if a substantial portion of the water column is removed in a given basin through evaporation, the loss of surface load will lead to an isostatic rebound of the asthenosphere, resulting in the shallowing of the affected basins (Figure 4. 5).

Figure 4 . 4

 44 Figure 4. 4 Schematic cross section of a well showing step-by-step paleo-bathymetric and sea level drawdown reconstruction

Figure 4 . 5

 45 Figure 4. 5 Schematic illustration of theoretical evolution of paleo water depth estimate (Y') as a function of the amplitude of the water level drawdown (XX') by evaporation showing the importance of the isostasy and water loss through evaporation components.

Chapter 4 -

 4 Late Messinian Mediterranean base level fall ___________________________________________________________________________________________ 178 4.3.3.1 Full basin scenario DSDP site 134 paleobathymetry at the end of stage 3 deposits accumulation for both scenarios can be derived from paleobathymetric reconstruction of the Western Mediterranean by Heida et al. (2021) (see the paper for detailed method). The uncertainty in the depth estimation related to lithological, flexural and thermal parameters is in the order of 100m (Heida et al., 2021). For sites 654 (Tyrrhenian Basin), 374 (Ionian Basin, Messina abyssal plain) and 376 (East

Chapter 4 -

 4 Late Messinian Mediterranean base level fall ___________________________________________________________________________________________ 179 by lateral variations in thicknesses, and the resolution of our results is within a range of several hundred meters. 4.3.3.2 Drawdown scenario Here we estimate the reduction in water column height (XX') required to obtain a water column thickness (X'-Y') at the end of the MSC in each well below the maximum reference depth inferred from the isotope measurement (X'Y'max in panel C and D of Figure 4.4). Considering possible rebound due to a water level drop, the paleobathymetry (Y') would have been shallower (see panel D in Figure 4.4) than estimated in the full basin scenario (Figure 4.4, panel B) during a low stand by a factor of one-third of the water-level change magnitude(=X'Y'full-X'Y'max), as the density contrast between air and the asthenosphere is three times larger than the density contrast between air and water (assuming densities if ~3000 kg/m 3 for the asthenosphere, 1000 kg/m 3 for water and 0 kg/m 3 for air). This means that the water level (X') during such drawdown, was higher than the required reduction in water column height suggests, when accounting for local isostatic compensation for the drawdown (in Figure 4.4 compare panel C and D). The procedure to estimate both the reduction in water column height and the corresponding water level is illustrated in Figure 4.4. 𝑷𝒂𝒏𝒆𝒍 𝑪: 𝑊𝑎𝑡𝑒𝑟 𝑑𝑒𝑝𝑡ℎ 𝑟𝑒𝑑𝑢𝑐𝑡𝑖𝑜𝑛 = 𝑋′𝑌′ ,-.. -𝑋′𝑌 O>P 𝑷𝒂𝒏𝒆𝒍 𝑫: 𝑋′ = 𝑊𝑎𝑡𝑒𝑟 𝑑𝑒𝑝𝑡ℎ 𝑟𝑒𝑑𝑢𝑐𝑡𝑖𝑜𝑛 *

Figure 4

 4 Figure 4. 6 EDX map of sample 654-30 (laminated gypsum). Areas containing both iron and sulfur are highlighted in orange and shown by the white arrows.

  (Figure 4. 7 and Figure 4. 8). The frequent occurrence of amorphous aggregates rich in iron and oxygen suggest the presence of iron oxides in these samples and supports the observation of Kuroda et al. (2016) for sites 374 and 376 (Figure 4. 7).Very rare larger micrometric framboidal pyrite crystals were found only in anhydrite samples and mudstone. In one sample of anhydrite pyrite is present as euhedral crystals with Fe/S ratio of 0.5.

Figure 4 . 7 Figure 4 . 8

 4748 Figure 4. 7 Secondary electron image (A) and (B) EDX map of sample 376-20C (gypsum siltite) showing the repartition of iron sulfide minerals (in orange). Note the abundance of iron oxides (in red) around the iron sulfides. (C) close up of the zone highlighted by a black square in (A) and (B) showing the nanometric size of the iron sulfides that seem to encrust other minerals. (D) EDX spectra of the zone in (C).

Figure 4 . 9

 49 Figure 4. 9 Concentration of iron sulfides within gypsum and marls interbeds of cores 134, 374, 376 and 654 as a function of their δ 34 S isotopic composition.

Chapter 4 -

 4 Late Messinian Mediterranean base level fall ___________________________________________________________________________________________ 187 gypsum they formed on, or close to, is an effectively infinite source of sulfate for microbial sulfate reducers(Hoareau et al., 2011).

Chapter 4 -

 4 Late Messinian Mediterranean base level fall ___________________________________________________________________________________________ 189

Chapter 4 -Figure 4 .

 44 Figure 4. 10 Theoretical evolution of sulfate (solid lines) and associated sulfide (dashed lines) during the consumption of a finite reservoir of sulfate by MSR (when fSO4 =1 no sulfate have been consumed, whereas fSO4 =1 represent the complete consumption of sulfate). Blue lines represent the respective values of sulfate and sulfide for a fractionation εMSR=-20‰ and red lines represent the evolution of sulfate-sulfide couple for the maximum fractionation possibly expressed during MSR.

Chapter 4 -

 4 Late Messinian Mediterranean base level fall ___________________________________________________________________________________________ 191 sulfurization, early diagenetic iron sulfides preserved in sedimentary rocks are most likely formed close to the sediment water interface (<1mbsf) and have a light isotopic composition (δ 34 S between -50 and -30‰).

Chapter 4 -

 4 Late Messinian Mediterranean base level fall __________________________________________________________________________________________ _ ∼350m) were primarily controlled by interglacial-glacial sea level oscillations, with pyrite δ 34 S displaying large amplitude variations (-46<δ 34 S<+20‰) and on average high values (δ 34 S=-21 ±15‰) (Pasquier et al., 2021a: see compilation in their supplementary materials in addition to the results of the paper). Conversely, the authors demonstrated that when the water depth of a basin was sufficiently large, no major sedimentary regime shifts were observed during glacial sea level drops and the pyrite preserved in the sediment s panning several glacialinterglacial oscillations had relatively invariant and 34 S-depleted values over time (δ 34 S=-51 ±5‰).

Chapter 4 -

 4 Late Messinian Mediterranean base level fall ___________________________________________________________________________________________ 194 4.5.2.3 Putting a water depth limit on the sedimentary control of iron sulfide isotopic composition Here we tried to find a minimal water column thickness (XY) at which sedimentary settings necessarily lead to the formation of iron sulfides that display highly depleted δ 34 S. The goal was to understand if the mechanistic relationship proposed in section 4.5.2.2 between basin depth (Y) and early formed iron sulfide isotopic composition is actually observed in the sedimentary record. This was done in order to understand if the large isotopic variations observed in iron sulfides preserved in stage 3 deposits could have been observed under their present-day deep water settings (between 1965mbsl and 4460 mbsl). We thus compiled 422 literature measurement of iron sulfide isotopic composition measured from sediments at less than 1m below sea floor (mbsf) in basins with various water column thicknesses (XY) (from 2 to 4974 mbsl) (orange dots in Figure 4. 11). We considered iron sulfides in sediments above 1mbsf to be representative of early diagenetic environments, and thus their isotopic composition to be comparable to the ones measured in our samples. To this compilation, we added 467 measurements of δ 34 S from the literature from sedimentary archives deposited in basins that didn't undergo significative vertical motion since sediment deposition, meaning that the depth of deposition (Y') can be assimilated to their actual depth (Y) (with depths from 150 to 3290mbsl) (blue dots in Figure 4. 11) (see Appendix B for the references, location and values used to build the compilation).

Chapter 4 -Figure 4 .

 44 Figure 4. 11 Iron sulfide isotopic composition as a function of the water depth of the sediments. Orange dots corresponds to iron sulfide isotopic data from the first meter of modern sediments from various locations (see table B.2 in the appendix section for full references and location of the samples) whereas purple dots correspond to iron sulfide from recent sedimentary archives (see table B.1 in the appendix section for full references and location of the samples). The green square represents the range of isotopic values measured in this study. Note that they span almost the entire range of values measured measured in shallow water sediments.

Chapter 4 -

 4 Figure 4. 10). First, marine basins exceeding 200m depth generally exhibit sulfate

Chapter 4 -

 4 Late Messinian Mediterranean base level fall ___________________________________________________________________________________________ 200 alternatively the large evidences supporting the partial desiccation of the Mediterranean during the MSC

4. 5 . 3 . 1

 531 Scenario 1: no drawdown To reconstruct the paleobathymetry Y' at each individual site, we first considered the constraints that were applied to each individual basin considering each basins filled to global Chapter 4-Late Messinian Mediterranean base level fall ___________________________________________________________________________________________ 201 sea-level (see section 4.3.3.1 for details on the approach). This was to assess if the sedimentary layers considered could have been deposited at depth <1500mbsl, without a significant sea level drawdown (XX') of the Mediterranean. If this hypothesis is verified, it means that the drastic decrease in water thicknesses (X'-Y') required to reach our X'Y'max maximum estimate have only been induced by post Messinian subsidence (YY'). Following the isostasy-based paleodepth reconstruction by Heida et al. (2021), the bathymetry Y' at site 134 at the end of stage 3 deposits accumulation would have been of approximately 2700m if no substantial base level fall of the western Mediterranean basin occurred. At site 654, assuming local isostasy, post-Messinian subsidence due to the sediment load would have reached 115 meters. In addition, a component of compaction of underlying sediments in the range of tens of meters, bring the Miocene-Pliocene boundary from its current depth of 2450 mbsl to about 2300 msbl. In the base of the site 654 well, continental conglomerates of unknown age are reported, overlain by a deepening-upwards sequence ranging from late Tortonian to Early Messinian. The sedimentary unit directly underlying the

Chapter 4 -

 4 Late Messinian Mediterranean base level fall ___________________________________________________________________________________________ 202

Chapter 4 -

 4 Late Messinian Mediterranean base level fall ___________________________________________________________________________________________ 203 and Anatolia, which started uplifting at least since the Middle Miocene (Symeou et al., 2018).

Chapter 4 -

 4 Late Messinian Mediterranean base level fall ___________________________________________________________________________________________ 206 processes in the subduction systems could have caused foundering of the basin, possibly explaining the exceptionally deep pinch-out of MSC evaporites on

Chapter 4 -

 4 Late Messinian Mediterranean base level fall ___________________________________________________________________________________________ 208 between the Western and the Eastern Mediterranean basins that could have evolved independently from each other during MSC stage 3 and possibly already during stage 2.

Chapter 5 -: 1 1

 51 Unraveling the roles of sulfide oxidation and of variable fractionation factors during microbial sulfate reduction ___________________________________________________________________________________________ 209 Chapter 5. Controls on iron sulfide multiple sulfur isotopic composition: unraveling the roles of variable fractionation factors during microbial sulfate reduction and of sulfide oxidation Guibourdenche, L. 1 and Aloisi, G. Université de Paris, Institut de physique du globe de Paris, CNRS, F-75005 Paris, France

Figure 5 . 1

 51 Figure 5. 1 Scheme representing how global ocean (blue) and restricted reservoir sulfur (yellow) cycle are linked. Sulfur bearing minerals generally precipitate in reservoirs that are partially or completely isolated from the open ocean virtually infinite sulfate reservoir. Blue arrows corresponds to the fluxes of sulfur controlling oceanic sulfate concentration on a long time scale whereas black arrows (amount of exchange with the oceanic sulfate reservoir and fluxes of sulfur species within the restricted sulfate reservoir) controls the isotopic composition of sedimentary sulfide preserved in the geological record.

Chapter 5 -:

 5 Unraveling the roles of sulfide oxidation and of variable fractionation factors during microbial sulfate reduction ___________________________________________________________________________________________ 211 timescale in which the size of the sulfur fluxes and isotopic composition change, are smaller

Chapter 5 -:Figure 5 . 2

 552 Figure 5. 2 Theoretical evolution of sulfate (solid lines) and associated sulfide (dashed lines) during the consumption of a finite reservoir of sulfate by microbial sulfate reduction (when fSO4 =1 no sulfate have been consumed, whereas fSO4 =1 represent the complete consumption of sulfate by MSR). Both the fractionation factor expressed during this biogeochemical process and the relative amount of sulfur consumed by it shape the isotopic composition of the sulfur species (sulfate and sulfide). Blue lines represent the respective values of sulfate and sulfidefor a fractionation εMSR=-20‰ and red lines represent the evolution of sulfate-sulfide couple for the maximum fractionation possibly expressed during MSR.

Chapter 5 -:

 5 Unraveling the roles of sulfide oxidation and of variable fractionation factors during microbial sulfate reduction ___________________________________________________________________________________________ 213 could lead to similar variations in the δ 34 SFeS isotopic composition of sulfide minerals as the ones observed when sulfide pooling occurs, adding further complexity in the interpretation of the sedimentary sulfur isotope signal. For example, in Figure5. 2, δ 34 SFeS>0 values can be attained if more than 90% of the sulfate has been consumed by MSR with a fractionation factor of -70 ‰, whereas only 20% of the sulfate needs to have been consumed by MSR to reach δ 34 SFeS>0 if the fractionation during MSR is equal to -20 ‰.

  the concurrent influence of sulfide pooling and variations in fractionation factors during MSR and (ii) constrains the occurrence of possible mixing or re-oxidative processes, we measured the multiple sulfur isotopic composition of iron sulfides and gypsum sulfate, and the oxygen isotopic composition of the same gypsum sulfate from 45 samples across a East/West transect of the Mediterranean, in the Messinian layers (both gypsum and marls interbeds) from DSDP and IODP cores 134, 374,376 and 654.Chapter 5-: Unraveling the roles of sulfide oxidation and of variable fractionation factors during microbial sulfate reduction ___________________________________________________________________________________________ 215

  Intuitively, the sulfate isotopic composition of open Mediterranean, deep-basin phase 3 gypsum deposits are expected to present a smaller isotopic offset from messinian oceanic sulfate than PLG from marginal basins because : (i) these deep basins possibly connected to the open Mediterranean during the MSC represent a larger reservoir of sulfate than marginal restricted shallow basins, and (ii) the biogeochemical cycling of sulfur is more intense in near shore zones of the ocean than in deep basins[START_REF] Bowles | Global rates of marine sulfate reduction and implications for sub-sea-floor metabolic activities[END_REF]. However, the few oxygen data existing from phase 3 gypsum deposits, from both ODP/DSDP cores (yellow;Blanc- Valleron et al. (1998);[START_REF] Fontes | Oxygen, Carbon, Sulfur and Hydrogen stable isotopes in carbonate and sulfate mineral phases of Neogene evaporites, sediments and in interstitial waters[END_REF];[START_REF] Pierre | Contribution ~ l'Jtude sgdimentologique et isotopique des evaporites messiniennes de la Mediterrane; implications geodynamiques[END_REF];Pierre and Fontes (1982);[START_REF] Pierre | Isotope Composition of Messinian sediments from the Mediterranean Sea as indicators of paleoenvironments and diagenesis[END_REF];Pierre and Rouchy (1990a);[START_REF] Ricchiuto | Stable Isotopic investigation of Messinian sulfate samples from DSDP. LEG 42. Eastern Mediterranean Sea[END_REF]) and more marginal basins in Sicily and Cyprus (green ;[START_REF] Aloisi | The geochemical riddle of "low-salinity gypsum" deposits[END_REF] García-Veigas et al. (2018);Pierre et al. (1982) indicates the contrary, with larger variations and deviation from MSC ocean (blue;Masterson et al. (2016);Waldeck et al. (

5 . 3

 53 Figure 5. 4 Lithological log of the cores studied in this work

  these species when transported(Choppala et al., 2017) (see Chapter 4 for discussion about this matter). Since MSR that produce the hydrogen sulfide necessary for iron sulfide formation generally occur in a zone where other oxidants have been consumed, iron sulfides measured in this study most Chapter 5-: Unraveling the roles of sulfide oxidation and of variable fractionation factors during microbial sulfate reduction ___________________________________________________________________________________________ 220 probably formed in an anoxic zone, at the bottom of the basin, possibly at or below the sediment water interface.

Figure 5 . 5

 55 Figure 5. 5 Multiple sulfur isotopic composition (∆ 33 SSO4 and δ 34 SSO4) of phase 3 calcium sulfate analysed in this study and existing data for PLG gypsum

Figure 5 . 9

 59 Figure 5. 9 Multiple sulfur isotopic composition (∆ 33 SFeS and δ 34 SFeS) of iron sulfide within phase 3 calcium sulfate and marls interbeds analysed in this study

Chapter 5 -

 5 : Unraveling the roles of sulfide oxidation and of variable fractionation factors during microbial sulfate reduction ___________________________________________________________________________________________ 229 transported (Choppala et al., 2017), and (iii) the cycle of hydration/dehydration of gypsum

Chapter 5 -

 5 : Unraveling the roles of sulfide oxidation and of variable fractionation factors during microbial sulfate reduction ___________________________________________________________________________________________ 230 scenarios to understand the controls on the isotopic compositions we observe. These scenarios represent the processes that occur in a range of realistic possible depositional conditions during the formation of iron sulfide.

Chapter 5 -

 5 : Unraveling the roles of sulfide oxidation and of variable fractionation factors during microbial sulfate reduction ___________________________________________________________________________________________ 231 expansion of the water-column oxygen minimum zone that is correlated to the isotopic composition of the sedimentary sulfides record(Pasquier et al., 2021b). This is because organic carbon availability is one of the main drivers controlling the rate at which sulfate is consumed by microbes. The authors interpret that during periods of increased organic matter delivery to the bottom of the basin, a more efficient consumption of porewater sulfate favored the production of 34 S-enriched hydrogen sulfide by distillation of isotopes, and this 34 S-enriched hydrogen sulfide was in turn preserved in the sediment record. During the Pleistocene in the Valle di Manche section in Calabria, it is the combination of both sedimentation rates and changes in organic matter delivery and sources, driven by transgressive-regressive cycles that were inferred to control the large variation in δ 34 Spyrite of sedimentary pyrites(Houghton et al., 2022). In any case, and regardless of the sedimentary control involved, these studies link large amplitude variations in the sulfur isotopic composition of sulfide to modulation in the distillation of sulfur isotopes in the sediments pore-water.

Chapter 5 -

 5 : Unraveling the roles of sulfide oxidation and of variable fractionation factors during microbial sulfate reduction ___________________________________________________________________________________________ 232 by diffusion, physical reworking or bioturbation of an undistilled sulfide formed at the top of the MSR zone a with a 34 S-enriched pooled sulfide formed deeper in the sediment (Pasquier et al., 2021a) (Figure 5. 10).

Figure 5 .

 5 Figure 5. 10 Structure of the numerical model simulating scenario 1 : the mixing between a pool of undistilled sulfide produced by the microbial reduction of the sulfate from an infinite reservoir and a pool of sulfide accumulated during the progressive consumption of porewater sulfate by MSR in a restricted reservoir. The result of the model gives the isotopic composition of the iron sulfide formed with negligible fractionation from this mix.

Chapter 5 -Figure 5 .

 55 Figure 5. 11 Measured values of the isotopic composition of iron sulfides preserved in calcium sulfate intervals and non-evaporitic interbeds from this study compared to modelled scenario 1. Individual curves corresponds to the modelled isotopic composition of sulfide for varying fSO4 (amount of sulfate from the restricted reservoir being consumed by MSR) and fmix (the amount of mixing between the pooled and unpooled reservoir of sulfide. In panel A ,εMSR=-70‰, in panel B εMSR=-50‰, and in panel εMSR=-20‰.

Figure 5 .

 5 Figure 5. 12 Concentration of iron sulfides within gypsum and marls interbeds of cores 134, 374, 376 and 654 as a function of their δ 34 S isotopic composition

Figure 5 .

 5 Figure 5. 13 Structure of the numerical model simulating scenario 2: the production of sulfide by MSR with variable fractionation factors in a restricted reservoir where sulfide accumulate during the progressive consumption of porewater sulfate by MSR. The result of the model gives the isotopic composition of the iron sulfide formed with negligible fractionation.

  Unraveling the roles of sulfide oxidation and of variable fractionation factors during microbial sulfate reduction ___________________________________________________________________________________________

  (B) from scenario 1, using equations (F) and (G), instead of fixed fractionation factors in equation (A) in order to model the combined effects of varying depletion of the sulfate reservoir by MSR consumption to MSR expressing the entire range of fractionation factor(εMSR) possible during this process.

Chapter 5 -Figure 5 .

 55 Figure 5. 14 Measured values of the isotopic composition of iron sulfides preserved in calcium sulfate intervals and non-evaporitic interbeds from this study compared to modelled scenario 2. Individual curves corresponds to the modelled isotopic composition of sulfide for varying fSO4 (amount of sulfate from the restricted reservoir being consumed by MSR) and εMSR the fractionation factor expressed during MSR.

Chapter 5 -

 5 : Unraveling the roles of sulfide oxidation and of variable fractionation factors during microbial sulfate reduction ___________________________________________________________________________________________ 243

___________________________________________________________________________________________ 244 Figure 5 .

 2445 Figure 5. 15 Structure of the numerical model simulating scenario 3: mixing between sulfide produced by MSR with variable fractionation factors with sulfide formed with the maximum fractionation possibly expressed (-70‰). The result of the model gives the isotopic composition of the iron sulfide formed with negligible fractionation.

Figure 5 .

 5 Figure 5. 16 Measured values of the isotopic composition of iron sulfides preserved in calcium sulfate intervals and non-evaporitic interbeds from this study compared to modelled scenario 3. Individual curves correspond to the modelled isotopic composition of sulfide resulting from the mixing of sulfide produced by MSR with varying fractionation factors εMSR.

  of our sulfide samples, varying fractionation factors are needed to explain the variability measured in our samples. This provide further evidence for the occurrence of varying fractionation factors in natural environment during MSR. How realistic this scenario is, and how to relate the variability in fractionation factors during MSR to environmental parameters shift, is nevertheless difficult to test because the eventual environmental controls involved in the expression of small fractionation factors during MSR other than organic matter availability are poorly constrained. 5.8.4 Scenario 4: Influence of re-oxidative processes In addition to MSR related processes, the occurrence of a re-oxidative cycling of sulfur during the formation of phase 3 deposits is suggested by the combined isotopic composition of δ 18 O and δ 34 S measured in sulfate from gypsum (Figure 5. 6 and Figure 5. 8) as previously proposed Chapter 5-: Unraveling the roles of sulfide oxidation and of variable fractionation factors during microbial sulfate reduction ___________________________________________________________________________________________ 247

  5Figure 5. 6) whereas the δ 18 O isotopic composition of the same sulfate present substantial deviation from messinian oceanic values (Figure 5. 8, Figure 5. 17).

Chapter 5 -Figure 5 .

 55 Figure 5. 17 Diagram showing the exisiting δ 34 S vs δ 18 O sulfate isotopic data for upper gypsum measured in marginal basins of the Mediterranean[START_REF] Aloisi | The geochemical riddle of "low-salinity gypsum" deposits[END_REF] García-Veigas et al., 2018; Pierre and Fontes, 1982) and those from deep basins of the Mediterranean measured in this study. The possible evolution of the isotopic composition of residual sulfate during MSR is indicated by the grey lines[START_REF] Fotherby | Modelling the Effects of Non-Steady State Transport Dynamics on the Sulfur and Oxygen Isotope Composition of Sulfate in Sedimentary Pore Fluids[END_REF].

  Unraveling the roles of sulfide oxidation and of variable fractionation factors during microbial sulfate reduction ___________________________________________________________________________________________ 249 had an isotopic composition close to Messinian oceanic value we consider that the calculated net fractionation factor was representing the closest estimate possible to reality. The εnet net fractionation factors for sulfides measured in this study mostly fall within the area defined by pure culture experiments of microbial sulfate reducers (Figure 5. 18).

Figure 5 .

 5 Figure 5. 18 Comparison between fractionation factors determined in laboratory experiments and net fractionation factors deduced from the multiple sulfur isotope measurement of iron sulfides from this study. Data from the literature used in this figure are the same than used in Chapter 3, see table 2.1..

Figure 5 .

 5 Figure 5. 19 Structure of the numerical model simulating scenario 3: mixing between sulfide produced by MSR with variable fractionation factors with sulfide formed with the maximum fractionation possibly expressed(-70‰). The result of the model gives the isotopic composition of the iron sulfide formed with negligible fractionation.

Figure 5 .

 5 phototrophic oxidation in pink Panel B,Figure5. 20), and chemolithotrophic oxidation in

Chapter 5 -

 5 : Unraveling the roles of sulfide oxidation and of variable fractionation factors during microbial sulfate reduction ___________________________________________________________________________________________ 254 re-oxidative processes possibly involved in the sedimentary environment, therefore the fractionation factors of a number of re-oxidative processes remains unknown. In the absence of fossils remains of sulfide oxidizers or measurement of reactive iron fraction in our samples, we have no other clue than sulfate-sulfide isotopic compositions to infer on the nature of the re-oxidative reaction.Finally, in this scenario, the isotopic variability measured in our samples can only be explained if a certain degree of variation in the fractionation factor ( 34 ƐMSR) was expressed during MSR (Figure5. 20). The extent of the required range of fractionation variation depends on the scenario considered with 34 ƐMSR varying between ∼-60 and -20‰ for the abiotic oxidation scenario (Panel A, Figure5. 20), ∼-60< 34 ƐMSR<-10‰ in the case of the phototrophic oxidation scenario (Panel B,Figure5. 20), and 34 ƐMSR oscillating of only 25‰ (between -65 and -40‰) if chemolithotrophic oxidation was the oxidative pathway allowing for the bulk of the H2S reoxidation (Panel C, Figure5. 20). From multiple sulfur isotopes only it is difficult to choose between scenario 3 (implying that only variable fractionation factors during MSR controlled iron sulfide isotopic composition) and scenario 4 (implying that both sulfide re-oxidation and variable fractionation factors shaped the isotopic composition of iron sulfides). However, the coupled δ 34 S -δ 18 O isotopic composition strongly suggest the occurrence of re-oxidative sulfur cycling. Together with the fact the almost complete re-oxidation of sulfide produced by MSR is understood to be common in marine sediments, this is in favor of scenario 4.

  This work suggests a tight coupling of microbially driven sulfate reduction and sulfide oxidation in 4 different large and deep basins of the Mediterranean at the end of the Messinian Salinity Crisis. By studying the very specific paleoenvironmental conditions of Chapter 5-: Unraveling the roles of sulfide oxidation and of variable fractionation factors during microbial sulfate reduction ___________________________________________________________________________________________ 255

Chapter 5 -

 5 : Unraveling the roles of sulfide oxidation and of variable fractionation factors during microbial sulfate reduction ___________________________________________________________________________________________ 256 basin of the Mediterranean in the Chapter 3 of this manuscript, this further demonstrate that this has been a common feature of the sulfur cycling in the entire Mediterranean Sea during

Figure 6 . 1

 61 Figure 6. 1 Dead Sea SaltGiant field trip, November 2021. Courtesy of Jimmy Moneron

Figure 6

 6 Figure 6. 2 A. Sulfur δ 34 S isotopic curve for both sulfate and sulfide in the PLG succession of the VdG basin, calibrated with the 100ka orbital eccentricity and precession curves[START_REF] Laskar | La2010: a new orbital solution for the long-term motion of the Earth[END_REF]. B. Sulfur δ 34 S isotopic composition of iron sulfide within marls and gypsum/anhydrite interbeds. Note the 34 Sdepleted isotopic composition of iron sulfides within marls (green crosses in panel A and orange whisker plot in panel B) compared to the overall higher δ 34 S of iron sulfides entrapped within calcium sulfate minerals (pink to purple dots in panel A and green whisker plotin panel B) in both A and B.

Figure 6 . 3

 63 Figure 6. 3 Modelled isotope composition of PLG deposits (stage 1) compared to measured values. Data points in (A) represent the Δ 33 S vs δ 34 S of sulfate and sulfide of gypsum and marl samples from the PLG unit of the VdG Basin. Data points in (B) represent the isotopic fractionation calculated for paired sulfate and sulfide in the various gypsum lithofacies (massive, banded, and branching selenite). Individual curves represent the modelled isotopic compositions in (A) and net fractionations factors in (B) for different values of fox (solid lines) and ƐMSR (dashed lines) the extent of isotopic fractionation during MSR. ƐMSR approaches -70‰ when thermodynamic equilibrium is reached and fox=1, indicating that all the sulfide produced by MSR is reoxidized. Note that in both diagrams most of the data fall in the area defined by fox>0.85, indicating the occurrence of nearly complete reoxidation of sulfide, and supports the occurrence of ongoing cryptic sulfur cycling during gypsum formation.

Chapter 6 -

 6 Conclusion and perspectives ___________________________________________________________________________________________ 264 precipitation in marginal basins of the Mediterranean during stage 1. This aspect is developed in Aloisi et al. (2022) and further details can be found in Appendix C of this manuscript (Figure 6. 5, panel C).

Figure 6 . 5

 65 Figure 6. 5 Schematic representation of the biogeochemical conditions that accompanied PLG deposition. (A) Deposition of organic rich marls during humid phases (precession minima). (B) Deposition of gypsum in the margins and marls in the depocenter during arid phases (precession maxima). The relative size of the arrows represents the approximate magnitude of fluxes. (C) Details of a gypsum twin.

Figure 6 . 6

 66 Figure 6. 6 Secondary electron image of sample 654-36B (marl) showing the nanometric size and shape of typical iron sulfide globules found in stage 3 deposits sampled in this study.

Chapter 6 -

 6 Conclusion and perspectives ___________________________________________________________________________________________ 266 diagenetic iron sulfides and their δ 34 S composition, with generally lower and less variable isotopic compositions in basins exceeding 1500m depth than in shallower basins (Figure 6. 7).

Figure 6 . 7

 67 Figure 6. 7 Early diagenetic iron sulfide isotopic composition as a function of the water depth of the sediments. Orange dots corresponds to iron sulfide isotopic data from the first meter of modern sediments from various locations (see tableB.2 in the appendix section for full references and location of the samples) whereas purple dots correspond to iron sulfide from recent sedimentary archives (see tableB.1 in the appendix section for full references and location of the samples). The green square represents the range of isotopic values measured in this study. Note that they span almost the entire range of values measured measured in shallow water sediments.

Chapter 6 -

 6 Conclusion and perspectives ___________________________________________________________________________________________ 268While the measurement of huge vertical variations in the rate of sulfide re-oxidation within sediments have recently highlighted the importance of accurate sulfide oxidation quantification for a more precise understanding of carbon cycling marine sediment[START_REF] Liu | Early diagenesis of iron and sulfur in Bornholm Basin sediments: The role of near-surface pyrite formation[END_REF][START_REF] Treude | Rapid sulfur cycling in sediments from the Peruvian oxygen minimum zone featuring simultaneous sulfate reduction and sulfide oxidation[END_REF], the horizontal (across various marine environments) and temporal variability in fluxes of sulfide re-oxidation as well as the variety of biotic and abiotic processes possibly involved are overlooked. This is important because the overall rate of sulfide reoxidation controls the amount of pyrite buried in the sediments. The quantification of long-term changes in the amount of sulfide being reoxidized therefore has implications for the understanding of the history of Earth oxygenation.

Chapter 6 -

 6 Conclusion and perspectives ___________________________________________________________________________________________ 269

Chapter 6 -

 6 Conclusion and perspectives ___________________________________________________________________________________________ 270 the role of a variable fractionation factor in controlling the isotopic composition of sedimentary sulfide. Although the exact triggers for the expression of a wide range of isotopic fractionation during MSR such as suggested in chapter 3 and 4 (see Figure 6. 3Figure 6. 4) are still poorly constrained, this work has allowed to distinguish the role of various fractionation factors expressed during MSR from the effect of Rayleigh distillation, that also has the potential to result in the production of large amplitude variations in the isotopic composition of sedimentary sulfide.Differentiating these two major controls on sulfide isotopic composition is not easy, but is nevertheless crucial, because in biogeochemical models of the sulfur cycle the burial fluxes of reduced sulfur depend on the MSR sulfur fractionation factor used(Kunzmann et al., 2017; Sansjofre et al., 2016; Tostevin et al., 2014) (Figure 6. 8).

Figure 6 . 8

 68 Figure 6. 8 Evolution of sulfate (solid lines) and associated sulfide (dashed lines) during the consumption of sulfate by microbial sulfate reduction (fSO4 =1 representing the complete consumption of sulfate by MSR). Blue lines represent the respective values of sulfate and sulfide for a fractionation εMSR=-20‰ and red lines represent the evolution of sulfate-sulfide couple for εMSR=-70‰

-

  Berner, R.A. (1984) Sedimentary pyrite formation: An update. Geochimica et Cosmochimica Acta 48, 605-615. -Berner, R.A. (1987) Models for carbon and sulfur cycles and atmospheric oxygen; application to Paleozoic geologic history. American Journal of Science 287, 177-196. -Berner, R.A. (1989) Biogeochemical cycles of carbon and sulfur and their effect on atmospheric oxygen over Phanerozoic time. Global and Planetary Change 1, 97-122. -Berner, R.A., Baldwin, T. and Holdren, G.R. (1979) Authigenic iron sulfides as paleosalinity indicators. Journal of Sedimentary Research 49, 1345-1350. -Bertoni, C. and Cartwright, J.A. (2005) 3D seismic analysis of circular evaporite dissolution structures, Eastern Mediterranean. Journal of the Geological Society 162, 909-926. -Bertran, E., Leavitt, W.D., Pellerin, A., Zane, G.M., Wall, J.D., Halevy, I., Wing, B.A. and Johnston, D.T. (2018) Deconstructing the Dissimilatory Sulfate Reduction Pathway: Isotope Fractionation of a Mutant Unable of Growth on Sulfate. Frontiers in Microbiology 9.-Bertran, E.,Waldeck, A., Wing, B.A., Halevy, I., Leavitt, W.D., Bradley, A.S. and Johnston, D.T. (2020) Oxygen isotope effects during microbial sulfate reduction: applications to sediment cell abundances. The ISME Journal 14, 1508-1519.-Bevington, P.R. and Robinson, D.K. (2003) Data reduction and error analysis.McGrawâ€"Hill, New York.

-

  Capella, W., Paleomagnetism and Paleomagnetism (2017) Gateway to a vanishing ocean: the tectono-sedimentary evolution of the Rifian Corridor (Morocco) and the Late Miocene isolation of the Mediterranean. UU Dept. of Earth Sciences.-[START_REF] Carnevale | Did the Mediterranean marine reflooding precede the Miocene/Pliocene boundary? Paleontological and geochemical evidence from upper Messinian sequences of Tuscany, Italy. Paleogeogr[END_REF] A leerfish (Teleostei, Carangidae) from the Messinian evaporite succession of the Vena del Gesso basin (Romagna Apennines, Italy): Palaeogeographical and palaeoecological implications.-Carnevale, G., Dela Pierre, F.,[START_REF] Carnevale | Fossil marine fishes and the "Lago Mare" event: Has the Mediterranean ever transformed into a brackish lake?[END_REF] Fossil marine fishes and the 'Lago Mare'event: Has the Mediterranean ever transformed into a brackish lake. Newsletters on Stratigraphy 51, 57-72.-Carnevale, G., Gennari, R., Lozar, F., Natalicchio, M., Pellegrino, L. and Dela Pierre, F. (2019) Living in a deep desiccated Mediterranean Sea: An overview of the Italian fossil record of the Messinian salinity crisis. -Carnevale, G. and Schwarzhans, W. (2022) Marine life in the Mediterranean during the Messinian salinity crisis: A paleoichthyological perspective. Rivista Italiana di Paleontologia e Stratigrafia 128, 283-324. -Carter, S.C., Paytan, A. and Griffith, E.M. (2020) Toward an Improved Understanding of the Marine Barium Cycle and the Application of Marine Barite as a Paleoproductivity Proxy. Minerals 10, 421. -Caruso, A., Blanc-Valleron, M.-M., Da Prato, S., Pierre, C. and Rouchy, J.M. (2020) The late Messinian "Lago-Mare" event and the Zanclean Reflooding in the Mediterranean Main text bibliography ___________________________________________________________________________________________ 280 Sea: New insights from the Cuevas del Almanzora section (Vera Basin, South-Eastern Spain). Earth-Science Reviews 200, 102993.

-

  Egger, M.,Riedinger, N., Mogollón, J.M. and Jørgensen, B.B. (2018) Global diffusive fluxes of methane in marine sediments. NatureGeoscience 11,[421][422][423][424][425] -El Euch-El Koundi, N., Ferry, S., Suc, J.-P., Clauzon, G., Melinte-Dobrinescu, M.,Gorini, C., Safra, A. and Zargouni, F. (2009) Messinian deposits and erosion in northern Tunisia: Inferences on Strait of Sicily during the Messinian Salinity Crisis. Terra Nova 21, 41-48. -El Kilany, A. (2018) Hydrologie et cycles biogéochimiques du soufre dans deux bassins marginaux de Méditerranée pendant la Crise de Salinité Messinienne, LOCEAN -Laboratoire d'Océanographie et du Climat : Expérimentations et Approches Numériques Université Sorbonne Paris Cité p. 274. -Eldridge, D., Kamyshny, A. and Farquhar, J. (2021) Theoretical estimates of equilibrium sulfur isotope effects among aqueous polysulfur and associated compounds with applications to authigenic pyrite formation and hydrothermal disproportionation reactions. Geochimica et Cosmochimica Acta. Eldridge, D.L. and Farquhar, J. (2018) Rates and multiple sulfur isotope fractionations associated with the oxidation of sulfide by oxygen in aqueous solution. Geochimica et Cosmochimica Acta 237, 240-260. -Eldridge, D.L., Guo, W. and Farquhar, J. (2016) Theoretical estimates of equilibrium sulfur isotope effects in aqueous sulfur systems: Highlighting the role of isomers in the sulfite and sulfoxylate systems. Geochimica et Cosmochimica Acta 195, 171-200. -Estrada, F., Ercilla, G., Gorini, C., Alonso, B., Vázquez, J.T., García-Castellanos, D., Juan, C., Maldonado, A., Ammar, A. and Elabbassi, M. (2011) Impact of pulsed Atlantic water inflow into the Alboran Basin at the time of the Zanclean flooding. Geo-Marine Letters 31, 361-376. -Evans, N.P., Turchyn, A.V., Gázquez, F., Bontognali, T.R., Chapman, H.J. and Hodell, D.A. (2015) Coupled measurements of δ18O and δD of hydration water and salinity of fluid inclusions in gypsum from the Messinian Yesares Member, Sorbas Basin (SE Spain). Earth and Planetary Science Letters 430, 499-510.-Faranda, C.,Gliozzi, E., Cipollari, P., Grossi, F., Darbaş, G., Gürbüz, K., Nazik, A., Gennari, R. and Cosentino, D. (2013) Messinian paleoenvironmental changes in the easternmost Mediterranean Basin: Adana Basin, southern Turkey. Turkish Journal of Earth Sciences 22, 839-863.-Farquhar, J., Johnston, D.T., Wing, B.A., Habicht, K.S., Canfield, D.E., Airieau, S. and Thiemens, M.H. (2003) Multiple sulphur isotopic interpretations of biosynthetic pathways: implications for biological signatures in the sulphur isotope record. Geobiology 1, 27-36. -Fike, D.A., Bradley, A.S. and Rose, C.V. (2015) Rethinking the ancient sulfur cycle. Annual Review of Earth and Planetary Sciences 43, 593-622. -Fike, D.A., Finke, N., Zha, J., Blake, G., Hoehler, T.M. and Orphan, V.J. (2009) The effect of sulfate concentration on (sub)millimeter-scale sulfide δ34S in hypersaline cyanobacterial mats over the diurnal cycle. Geochimica et Cosmochimica Acta 73, 6187-6204. -Fike, D.A., Gammon, C.L., Ziebis, W. and Orphan, V.J. (2008) Micron-scale mapping of sulfur cycling across the oxycline of a cyanobacterial mat: a paired nanoSIMS and CARD-FISH approach. The ISME Journal 2, 749-759.

  ) A shallow water record of the onset of the Messinian salinity crisis in the Adriatic foredeep (Legnagnone section, Northern Apennines). Palaeogeography, Palaeoclimatology, Palaeoecology 386, 145-164. -Ghosh, P. and Brand, W.A. (2003) Stable isotope ratio mass spectrometry in global climate change research. International Journal of Mass Spectrometry 228, 1-33. -Gilhooly, W.P., Reinhard, C.T. and Lyons, T.W. (2016) A comprehensive sulfur and oxygen isotope study of sulfur cycling in a shallow, hyper-euxinic meromictic lake. Geochimica et Cosmochimica Acta 189, 1-23. -Gilly, W.F., Beman, J.M., Litvin, S.Y. and Robison, B.H. (2013) Oceanographic and Biological Effects of Shoaling of the Oxygen Minimum Zone. Annual Review of Marine Science 5, 393-420.

-

  ) Sulphur geodynamic cycle. Scientific Reports 5, 8330. -Kah, L.C., Lyons, T.W. and Chesley, J.T. (2001) Geochemistry of a 1.2 Ga carbonateevaporite succession, northern Baffin and Bylot Islands: implications for Mesoproterozoic marine evolution. Precambrian Research 111, 203-234. -Kao, S.-J., Horng, C.-S., Roberts, A.P. and Liu, K.-K. (2004) Carbon-sulfur-iron relationships in sedimentary rocks from southwestern Taiwan: influence of geochemical environment on greigite and pyrrhotite formation. Chemical Geology 203, 153-168. -Kaplan, I.R. and Hulston, J.R. (1966) The isotopic abundance and content of sulfur in meteorites. Geochimica et Cosmochimica Acta 30, 479-496. -Karakitsios, V., Roveri, M., Lugli, S., Manzi, V., Gennari, R., Antonarakou, A., Triantaphyllou, M., Agiadi, K., Kontakiotis, G., Kafousia, N. and de Rafelis, M. (2017) A record of the Messinian salinity crisis in the eastern Ionian tectonically active domain (Greece, eastern Mediterranean). Basin Research 29, 203-233. -Kastens, K. and Mascle, J. (1990) The geological evolution of the Tyrrhenian Sea: An introduction to the scientific results of ODP Leg 107, Proceedings of the Ocean Drilling Program, Scientific Results. College Station, TX (Ocean Drilling Program), p. 26. Kastens, K., Mascle, J. and Auroux, C. (1987) Site 654: Upper sardinian margin. Kastens, KA, et al., Proceedings of the Ocean Drilling Program, Initial reports 107, 747-875.

  ) Metastable Iron (Mono)sulfides in the Shallow-Sea Hydrothermal Sediments of Milos, Greece. ACS Earth and Space Chemistry 6, 920-931. -Kouwenhoven, T., Seidenkrantz, M.-S. and Van der Zwaan, G. (1999) Deep-water changes: the near-synchronous disappearance of a group of benthic foraminifera from the Late Miocene Mediterranean. Palaeogeography, Palaeoclimatology, Palaeoecology 152, 259-281. -Kouwenhoven, T.J., Hilgen, F.J. and van der Zwaan, G.J. (2003) Late Tortonian-early Messinian stepwise disruption of the Mediterranean-Atlantic connections: constraints from benthic foraminiferal and geochemical data. Palaeogeography, Palaeoclimatology, Palaeoecology 198, 303-319. -Kouwenhoven, T.J., Morigi, C., Negri, A., Giunta, S., Krijgsman, W. and Rouchy, J.M. (2006) Paleoenvironmental evolution of the eastern Mediterranean during the Messinian: Constraints from integrated microfossil data of the Pissouri Basin (Cyprus). Marine Micropaleontology 60, 17-44. -Kraal, P., Burton, E.D. and Bush, R.T. (2013) Iron monosulfide accumulation and pyrite formation in eutrophic estuarine sediments. Geochimica et Cosmochimica Acta 122, 75-88. -Krijgsman, W., Capella, W., Simon, D., Hilgen, F.J., Kouwenhoven, T.J., Meijer, P.T., Sierro, F.J., Tulbure, M.A., van den Berg, B.C. and van der Schee, M. (2018) The Main text bibliography ___________________________________________________________________________________________ 291 Gibraltar corridor: Watergate of the Messinian salinity crisis. Marine Geology 403, 238-246.-Krijgsman, W., Hilgen, F.J., Raffi, I., Sierro, F.J. and Wilson, D.S. (1999) Chronology, causes and progression of the Messinian salinity crisis. Nature 400, 652-655.-Kristensen, E., Bodenbender, J., Jensen, M.H., Rennenberg, H. and Jensen, K.M. (2000) Sulfur cycling of intertidal Wadden Sea sediments (Konigshafen, Island of Sylt, Germany): sulfate reduction and sulfur gas emission. Journal of Sea Research 43, 93-104. -Kunzmann, M., Bui, T.H., Crockford, P.W., Halverson, G.P., Scott, C., Lyons, T.W. and Wing, B.A. (2017) Bacterial sulfur disproportionation constrains timing of Neoproterozoic oxygenation. Geology 45, 207-210.

  ) Determination of multiple sulfur isotopes in glasses: A reappraisal of the MORB δ34S. Chemical Geology 334, 189-198. -Laskar, J., Fienga, A., Gastineau, M. and Manche, H. (2011) La2010: a new orbital solution for the long-term motion of the Earth. Astronomy & Astrophysics 532, A89. -Leavitt, W.D., Halevy, I., Bradley, A.S. and Johnston, D.T. (2013) Influence of sulfate reduction rates on the Phanerozoic sulfur isotope record. Proceedings of the National Academy of Sciences 110, 11244.-Lécuyer, C. (2016) Seawater residence times of some elements of geochemical interest and the salinity of the oceans. Bulletin de la Société Géologique de France 187, 245-260.

  ) Miocene Atlantic-Mediterranean seaways in the Betic Cordillera (southern Spain). Revista de la Sociedad Geológica de España 27, 175-186.-Marzocchi, A., Flecker, R., van Baak, C.G.C., Lunt, D.J. and Krijgsman, W. (2016) Mediterranean outflow pump: An alternative mechanism for the Lago-mare and the end of the Messinian Salinity Crisis. Geology 44, 523-526.-Masterson, A., Alperin, M.J., Berelson, W.M. and Johnston, D.T. (2018) Interpreting multiple sulfur isotope signals in modern anoxic sediments using a full diagenetic model (California-Mexico margin: Alfonso Basin). American Journal of Science 318, 459-490. -Masterson, A.L., Wing, B.A., Paytan, A., Farquhar, J. and Johnston, D.T. (2016) The minor sulfur isotope composition of Cretaceous and Cenozoic seawater sulfate. Paleoceanography 31, 779-788. -Mayer-Eymar, K. (1867) Catalogue systématique et descriptif des fossiles des terrains tertiaires qui se trouvent au Musée féderal de Zürich. Librairie Schabelitz. -McIver, R.D. and RD, M. (1973) GEOCHEMICAL SIGNIFICANCE OF GAS AND GASOLINE-RANGE HYDROCARBONS AND OTHER ORGANIC MATTER IN A MIOCENE SAMPLE FROM SITE 134-BALEARIC ABYSSAL PLAIN.

--

  (1997) Empirical relationships for use in global diagenetic models. Deep Sea Research Part I: Oceanographic Research Papers 44, 327-344.-Mills, J.V.,Antler, G. and Turchyn, A.V. (2016) Geochemical evidence for cryptic sulfur cycling in salt marsh sediments. Earth and Planetary Science Letters 453, 23-32.-Mills, J.V.,Gomes, M.L., Kristall, B., Sageman, B.B., Jacobson, A.D. and Hurtgen, M.T. (2017) Massive volcanism, evaporite deposition, and the chemical evolution of the Early Cretaceous ocean. Geology 45, 475-478.-Mocnik, A.,[START_REF] Mocnik | The Messinian Salinity Crisis in the West-Mediterranean Basins: comparison between two rifted margins[END_REF] The Messinian Salinity Crisis in the West-Mediterranean Basins: comparison between two rifted margins, GNGTS 2014. Mosetti Tecniche Grafiche, Trieste, pp. 100-"-".-Morse, J.W. (1991) Oxidation kinetics of sedimentary pyrite in seawater. Geochimica et Cosmochimica Acta 55, 3665-3667. -Muller, E. (2016) Déconvolution du cycle du soufre à l'Archéen par l'analyse in situ des 4 isotopes du soufre, Geochimie des Isotopes Stables Université de Paris, Institut de Physique du Globe de Paris. Natalicchio, M., Birgel, D., Dela Pierre, F., Ziegenbalg, S., Hoffmann-Sell, L., Gier, S. and Peckmann, J. (2022) Messinian bottom-grown selenitic gypsum: An archive of microbial life. Geobiology. -Natalicchio, M., Dela Pierre, F., Birgel, D., Brumsack, H., Carnevale, G., Gennari, R., Gier, S., Lozar, F., Pellegrino, L., Sabino, M., Schnetger, B. and Peckmann, J. (2019) Paleoenvironmental change in a precession-paced succession across the onset of the Messinian salinity crisis: Insight from element geochemistry and molecular fossils. Palaeogeography, Palaeoclimatology, Palaeoecology 518, 45-61. -Natalicchio, M., Dela Pierre, F., Lugli, S., Lowenstein, T., Feiner, S., Ferrando, S., Manzi, V., Roveri, M. and Clari, P. (2014) Did Late Miocene (Messinian) gypsum precipitate from evaporated marine brines? Insights from the Piedmont Basin (Italy). Geology. -Natalicchio, M., Pellegrino, L., Clari, P., Pastero, L. and Dela Pierre, F. (2021) Gypsum lithofacies and stratigraphic architecture of a Messinian marginal basin (Piedmont Basin, NW Italy). Sedimentary Geology 425, 106009. -Nesteroff, W. (1973) Mineralogy, petrography, distribution, and origin of the Messinian Mediterranean evaporites. Initial reports of the deep sea drilling project 13, 673-693. -Nichols, G., Williams, E. and Paola, C. (2009) Sedimentary processes, environments and basins: a tribute to Peter Friend. John Wiley & Sons. -Ochoa, D., Sierro, F.J., Lofi, J., Maillard, A., Flores, J.-A. and Suárez, M. (2015) Synchronous onset of the Messinian evaporite precipitation: First Mediterranean offshore evidence. Earth and Planetary Science Letters 427, 112-124.-Ogniben, L. (1957) Secondary gypsum of the Sulphur Series, Sicily, and the so-called integration. Journal of Sedimentary Research 27, 64-79.-Ono, S.,Wing, B., Johnston, D., Farquhar, J. and Rumble, D. (2006) Mass-dependent fractionation of quadruple stable sulfur isotope system as a new tracer of sulfur biogeochemical cycles. Geochimica et Cosmochimica Acta 70, 2238-2252.-Orszag-Sperber, F., Caruso, A., Blanc-Valleron, M.-M., Merle, D. and Rouchy, J.M. (2009) The onset of the Messinian salinity crisis: Insights from Cyprus sections. Sedimentary Geology 217, 52-64. -Ortí, F., Rosell, L. and Anadón, P. (2010) Diagenetic gypsum related to sulfur deposits in evaporites (Libros Gypsum, Miocene, NE Spain). Sedimentary Geology 228, 304-318. -Palcu, D., Patina, I., Sandric, I., Lazarev, S., Vasiliev, I., Stoica, M. and Krijgsman, W. (2021) Late Miocene megalake regressions in Eurasia. Scientific Reports 11. -Panieri, G., Lugli, S., Manzi, V., Palinska, K.A. and Roveri, M. (2008) Microbial communities in Messinian evaporite deposits of the Vena del Gesso (northern Apennines, Italy). Stratigraphy 5, 343-352. -Panieri, G., Lugli, S., Manzi, V., Roveri, M., Schreiber, B.C. and Palinska, K.A. (2010) Ribosomal RNA gene fragments from fossilized cyanobacteria identified in primary gypsum from the late Miocene, Italy. Geobiology 8, 101-111. Pashko, P., Milushi, I. and Hoxha, V. (2019) THE MESSINIAN EVAPORITES OF THE PREADRIATIC FORELAND BASIN (ALBANIA). -Pasquier, V., Bryant, R.N., Fike, D.A. and Halevy, I. (2021a) Strong local, not global, controls on marine pyrite sulfur isotopes. Science Advances 7, eabb7403. -Pasquier, V., Fike, D.A. and Halevy, I. (2021b) Sedimentary pyrite sulfur isotopes track the local dynamics of the Peruvian oxygen minimum zone. Nature Communications 12, 4403.

--

  Pierre, C. and Rouchy, J.M. (1990) 13. SEDIMENTARY AND DIAGENETIC EVOLUTION OF MESSINIAN EVAPORITES IN THE TYRRHENIAN SEA (ODP LEG 107, SITES 652, 653, AND 654): PETROGRAPHIC, MINERALOGICAL, AND STABLE ISOTOPE RECORDS1. -Pinardi, N. and Masetti, E. (2000) Variability of the large scale general circulation of the Mediterranean Sea from observations and modelling: a review. Palaeogeography, Palaeoclimatology, Palaeoecology 158, 153-173. -Popescu, S.-M., Dalibard, M., Suc, J.-P., Barhoun, N., Melinte-Dobrinescu, M.-C., Bassetti, M.A., Deaconu, F., Head, M.J., Gorini, C., Do Couto, D., Rubino, J.-L., Auxietre, J.-L. and Floodpage, J. (2015) Lago Mare episodes around the Messinian-Zanclean boundary in the deep southwestern Mediterranean. Marine and Petroleum Geology 66, 55-70. -Potter, P.E. and Szatmari, P. (2009) Global Miocene tectonics and the modern world. Earth-Science Reviews 96, 279-295. -Prince, J.K.G., Rainbird, R.H. and Wing, B.A. (2019) Evaporite deposition in the mid-Neoproterozoic as a driver for changes in seawater chemistry and the biogeochemical cycle of sulfur. Geology 47, 375-379. -Pruden, G. and Bloomfield, C. (1968) The determination of iron(II) sulphide in soil in the presence of iron(III) oxide. Analyst 93, 532-534. -Raab, M., Friedman, G.M., Spiro, B., Starinsky, A. and Zak, I. (1997) The geological history of Messinian (upper miocene) evaporites in the Central Jordan Valley (Israel) Ryan, W., Hsu, K., Cita, M., Dumitrica, P., Lort, J., Maync, W., Nesterhoff, W., Pautot, G., Stradner, H. and Wezel, F. (1973) Boundary of Sardinia Slope with Balearic abyssal Plain-Site 133 and 134. Initial Reports of the Deep Sea Drilling Project 13, 465-514. -Ryan, W.B. (1976) Quantitative evaluation of the depth of the western Mediterranean before, during and after the Late Miocene salinity crisis. Sedimentology 23, 791-813. -Ryan, W.B. (2009) Decoding the Mediterranean salinity crisis. Sedimentology 56, 95-136. -Ryan, W.B.F. and Cita, M.B. (1978) The nature and distribution of Messinian erosional surfaces -Indicators of a several-kilometer-deep Mediterranean in the Miocene. Marine Geology 27, 193-230. -Sabino, M., Birgel, D., Natalicchio, M., Dela Pierre, F. and Peckmann, J. (2022) Carbon isotope excursions during the late Miocene recorded by lipids of marine Thaumarchaeota, Piedmont Basin, Mediterranean Sea. Geology 49.

  [START_REF] Sabino | Climatic and hydrologic variability in the northern Mediterranean across the onset of the Messinian salinity crisis[END_REF] Climatic and hydrologic variability in the northern Mediterranean across the onset of the Messinian salinity crisis. Palaeogeography, Palaeoclimatology, Palaeoecology 545, 109632. -Saitoh, M., Ueno, Y., Isozaki, Y. and Yoshida, N. (2021) Multiple sulfur isotope chemostratigraphy across the Permian-Triassic boundary at Chaotian, China: Implications for a shoaling model of toxic deep-waters. Island Arc 30, e12398.

--

  [START_REF] Topper | Precessional control of Sr ratios in marginal basins during the Messinian Salinity Crisis?[END_REF] Precessional control of Sr ratios in marginal basins during the Messinian Salinity Crisis. Geochemistry, Geophysics, Geosystems.-Topper, R.P.M., Flecker, R., Meijer, P.T. and Wortel, M.J.R. (2011) A box model of the Late Miocene Mediterranean Sea: Implications from combined 87Sr/86Sr and salinity data. Paleoceanography 26. -Topper, R.P.M. and Meijer, P.T. (2015a) Changes in Mediterranean circulation and water characteristics due to restriction of the Atlantic connection: a high-resolution ocean model. Clim. Past 11, 233-251. -Topper, R.P.M. and Meijer, P.T. (2015b) The precessional phase lag of Messinian gypsum deposition in Mediterranean marginal basins. Palaeogeography, Palaeoclimatology, Palaeoecology 417, 6-16.-Torfstein, A. and Turchyn, A.V. (2017) Rates and Cycles of Microbial Sulfate Reduction in the Hyper-Saline Dead Sea over the Last 200 kyrs from Sedimentary δ34S and δ18O(SO4). Frontiers in Earth Science 5. -Torres, M.A., West, A.J. and Li, G. (2014) Sulphide oxidation and carbonate dissolution as a source of CO2 over geological timescales. Nature 507, 346-349. Tostevin, R., Turchyn, A.V., Farquhar, J., Johnston, D.T., Eldridge, D.L., Bishop, J.K.B. and McIlvin, M. (2014) Multiple sulfur isotope constraints on the modern sulfur cycle. Earth and Planetary Science Letters 396, 14-21. -Treude, T., Hamdan, L.J., Lemieux, S., Dale, A.W. and Sommer, S. (2021) Rapid sulfur cycling in sediments from the Peruvian oxygen minimum zone featuring simultaneous sulfate reduction and sulfide oxidation. Limnology and Oceanography n/a. -Treude, T., Niggemann, J., Kallmeyer, J., Wintersteller, P., Schubert, C.J., Boetius, A. and Jørgensen, B.B. (2005) Anaerobic oxidation of methane and sulfate reduction along the Chilean continental margin. Geochimica et Cosmochimica Acta 69, 2767-2779. -Tudge, A.P. and Thode, H.G. (1950) THERMODYNAMIC PROPERTIES OF ISOTOPIC COMPOUNDS OF SULPHUR. Canadian Journal of Research 28b, 567-578. -Turchyn, A., Sivan, O. and Schrag, D. (2006a) Oxygen isotopic composition of sulfate in deep sea pore fluid: evidence for rapid sulfur cycling. Geobiology 4, 191-201. -Turchyn, A., Sivan, O. and Schrag, D.P. (2006b) Oxygen isotopic composition of sulfate in deep sea pore fluid: evidence for rapid sulfur cycling. Geobiology 4, 191-201. -Turchyn, A.V. and Schrag, D.P. (2004) Oxygen Isotope Constraints on the Sulfur Cycle over the Past 10 Million Years. Science 303, 2004. -Turchyn, A.V. and Schrag, D.P. (2006) Cenozoic evolution of the sulfur cycle: Insight from oxygen isotopes in marine sulfate. Earth and Planetary Science Letters 241, 763-779. -Tzevahirtzian, A., Caruso, A., Scopelliti, G., Baudin, F. and Blanc-Valleron, M.-M. (2022) Onset of the Messinian Salinity Crisis: Sedimentological, petrographic and geochemical characterization of the pre-salt sediments from a new core (Caltanissetta Basin, Sicily). Marine and Petroleum Geology, 105686. -Usiglio, M. (1849) Études sur la composition de l'eau de la Méditerranée et sur l'exploitation des sels qu'elle contient. Annales Chim. Phys., Ser 3 27, 172-191. -Utrilla, R., Pierre, C., Orti, F. and Pueyo, J.J. (1992) Oxygen and sulphur isotope compositions as indicators of the origin of Mesozoic and Cenozoic evaporites from Spain. Chemical Geology 102, 229-244. -Vai, G.B. and Ricci-Lucchi, F. (1977) Algal crusts, autochthonous and clastic gypsum in a cannibalistic evaporite basin: a case history from the Messinian of Northern Apennines. Sedimentology 24, 211-244. -Van Couvering, J., A., Castradori, D., Cita, M.B., J. Hilgen, F. and Rio, D. (2000) The base of the Zanclean Stage and of the Pliocene Series. International Union of Geological Sciences 23, 179-187. -Van Driessche, A., Canals, A., Ossorio, M., Reyes, R.C. and García-Ruiz, J.M. (2016) Unraveling the Sulfate Sources of (Giant) Gypsum Crystals Using Gypsum Isotope Fractionation Factors. The Journal of Geology 124, 235-245. van Santvoort, P.J.M., de Lange, G.J., Thomson, J., Cussen, H., Wilson, T.R.S., Krom, M.D. and Ströhle, K. (1996) Active post-depositional oxidation of the most recent sapropel (S1) in sediments of the eastern Mediterranean Sea. Geochimica et Cosmochimica Acta 60, 4007-4024. -van Vliet, D.M., von Meijenfeldt, F.B., Dutilh, B.E., Villanueva, L., Sinninghe Damsté, J.S., Stams, A.J. and Sánchez-Andrea, I. (2021) The bacterial sulfur cycle in expanding dysoxic and euxinic marine waters. Environmental microbiology 23, 2834-2857. -Vasiliev, I., Karakitsios, V., Bouloubassi, I., Agiadi, K., Kontakiotis, G., Antonarakou, A., Triantaphyllou, M., Gogou, A., Kafousia, N., Rafélis, M., Zarkogiannis, S., Kaczmar, F., Parinos, C. and Pasadakis, N. (2019) Large Sea Surface Temperature, Salinity, and Productivity-Preservation Changes Preceding the Onset of the Messinian Salinity Crisis in the Eastern Mediterranean Sea. Paleoceanography and Paleoclimatology 34. -Volvoikar, S., Mazumdar, A., Peketi, A., Dewangan, P., Sawant, B., Manaskanya, A., Goswami, H., Das, D. and Pujari, S. (2020) Contrasting sulfidization in the turbidite and hemipelagic sediments of Bengal Fan. Marine and Petroleum Geology 118, 104408. -Waldeck, A.R., Olson, H.C., Yao, W., Blättler, C.L., Paytan, A., Hodell, D.A. and Johnston, D.T. (2022) Calibrating the triple oxygen isotope composition of evaporite minerals as a proxy for marine sulfate. Earth and Planetary Science Letters 578, 117320. -Warren, J.K. (2006) Evaporites: sediments, resources and hydrocarbons. Springer Science & Business Media.-Warren, J.K. (2010) Evaporites through time: Tectonic, climatic and eustatic controls in marine and nonmarine deposits. Earth-Science Reviews 98, 217-268.-Wasmund, K., Mußmann, M. and Loy, A. (2017) The life sulfuric: microbial ecology of sulfur cycling in marine sediments. Environ Microbiol Rep 9, 323-344.

-

  [START_REF] Wing | Intracellular metabolite levels shape sulfur isotope fractionation during microbial sulfate respiration[END_REF] Intracellular metabolite levels shape sulfur isotope fractionation during microbial sulfate respiration. Proceedings of the National Academy of Sciences of the United States of America 111, 18116-18125. -Wortmann, U.G. and Chernyavsky, B.M. (2007) Effect of evaporite deposition on Early Cretaceous carbon and sulphur cycling. Nature 446, 654-656. -Wortmann, U.G. and Paytan, A. (2012) Rapid Variability of Seawater Chemistry Over the Past 130 Million Years. Science 337, 334-336. -Wu, N., Farquhar, J., Strauss, H., Kim, S.-T. and Canfield, D. (2010) Evaluating the Sisotope fractionation associated with Phanerozoic pyrite burial. Geochimica et Cosmochimica Acta 74. -Yao, W., Paytan, A. and Wortmann, U.G. (2018) Large-scale ocean deoxygenation during the Paleocene-Eocene Thermal Maximum. Science 361, 804-806. Young, E.D., Galy, A. and Nagahara, H. (2002) Kinetic and equilibrium mass-dependent isotope fractionation laws in nature and their geochemical and cosmochemical significance. Geochimica et Cosmochimica Acta 66, 1095-1104. -Zachariasse, W.J., Kontakiotis, G., Lourens, L.J. and Antonarakou, A. (2021) The Messinian of Agios Myron (Crete, Greece): A key to better understanding of diatomite formation on Gavdos (south of Crete). Palaeogeography, Palaeoclimatology, Palaeoecology 581, 110633. -Zeebe, R.E. (2010) A new value for the stable oxygen isotope fractionation between dissolved sulfate ion and water. Geochimica et Cosmochimica Acta 74, 818-828. -Zerkle, A.L., Farquhar, J., Johnston, D.T., Cox, R.P. and Canfield, D.E. (2009) Fractionation of multiple sulfur isotopes during phototrophic oxidation of sulfide and elemental sulfur by a green sulfur bacterium. Geochimica et Cosmochimica Acta 73, 291-306. -Zerkle, A.L., Jones, D.S., Farquhar, J. and Macalady, J.L. (2016) Sulfur isotope values in the sulfidic Frasassi cave system, central Italy: A case study of a chemolithotrophic Sbased ecosystem. Geochimica et Cosmochimica Acta 173, 373-386. -Zhang, G., Zhang, X., Hu, D., Li, D., Algeo, T.J., Farquhar, J., Henderson, C.M., Qin, L., Shen, M., Shen, D., Schoepfer, S.D., Chen, K. and Shen, Y. (2017) Redox chemistry changes in the Panthalassic Ocean linked to the end-Permian mass extinction and delayed Early Triassic biotic recovery. Proc Natl Acad Sci U S A 114, 1806-1810. -Zhang, J.-Z. and Millero, F.J. (1993) The products from the oxidation of H2S in seawater. Geochimica et Cosmochimica Acta 57, 1705-1718.

Figure B. 1

 1 Figure B. 1 Map of the global ocean bathymetry. Each dot or square represent a site included in our compilation. Note the bias of this compilation toward coastal region of the ocean/or marginal seas. The color of the dots depend on the most enriched δ 34 S data at each site included in our compilations (see figure to see the significance of each color). Map realized with http://www.geomapapp.org

Fig. 1 .

 1 Fig. 1. Marginal basin gypsum deposits formed during the Messinian salinity crisis. (a) Map of Mediterranean Sea showing the modern distribution of offshore and onshore deposits related to the Messinian Salinity Crisis. The geographical location of marginal basins (Sorbas, Piedmont, Vena del Gesso, Caltanissetta, Polemi and Mesaoria basins and the Catanzaro Trough) is shown, as well as the location of the Vaccare `s Lagoon (Camargue Area, south of France), located 25 km away from the Aigues Mortes salt works; (b) Stratigraphic sections investigated in the present work.

  Fig.2. Geochemical evidence for the existence of low-salinity evaporitic gypsum deposits in the Mediterranean Sea. Vertical bars in panel (a) represent the frequency (left vertical axis) of salinity measurements made in primary fluid inclusion from gypsum from the Piedmont and Vena del Gesso basins (northern Italy) (this work), the Catanzaro Trough (Calabria, Southern Italy)[START_REF] Costanzo | Messinian twinned selenite from the Catanzaro Trough, Calabria, Southern Italy: field, petrographic and fluid inclusion perspectives[END_REF] and the Sorbas basin[START_REF] Evans | Coupled measurements of delta O-18 and delta D of hydration water and salinity of fluid inclusions in gypsum from the Messinian Yesares Member, Sorbas Basin (SE Spain)[END_REF]; blue circles represent the salinity-d 18 O composition of Sorbas basin gypsum fluid inclusions[START_REF] Evans | Coupled measurements of delta O-18 and delta D of hydration water and salinity of fluid inclusions in gypsum from the Messinian Yesares Member, Sorbas Basin (SE Spain)[END_REF]. Coloured circles in panels (b) to (e) are the isotope composition (d 18 O H2O and dD H2O ) of gypsum parent waters from the Caltanissetta (Sicily, Italy), Vena del Gesso (Apennines, Italy), Piedmont (Italy) (this work) and Sorbas (Spain)[START_REF] Evans | Coupled measurements of delta O-18 and delta D of hydration water and salinity of fluid inclusions in gypsum from the Messinian Yesares Member, Sorbas Basin (SE Spain)[END_REF]. Crosses, and the vertical grey band in panel f, indicate the isotope composition of parent waters of gypsum formed in modern salt works of the Mediterranean region[START_REF] Fontes | Fractionnement isotopique dans l'eau de cristallisation du sulfate de calcium[END_REF][START_REF] Fontes | Fractionnement isotopique de l'hydroge `ne dans l'eau de cristallisation du gypse[END_REF] Longinelli, 1979;[START_REF] Evans | Coupled measurements of delta O-18 and delta D of hydration water and salinity of fluid inclusions in gypsum from the Messinian Yesares Member, Sorbas Basin (SE Spain)[END_REF]. The black dot is the isotope composition of surface Atlantic water in from of Gibraltar(Benetti et al., 2017a). The continuous green line, followed by the green dotted line in panels (b) to (f), represent the model evolution in the d 18 O H2O -dD H2O space during the evaporation of Mediterranean Sea water under evaporative conditions based on in-situ atmospheric measurements from the Vaccare `s Lagoon (Camargue area, southern France)[START_REF] Delattre | Deuterium excess in the atmospheric water vapour of a Mediterranean coastal wetland: regional vs. local signatures[END_REF]. The green dot with a black circle separating the continuous green line from the green dotted line represents the gypsum saturation point; the d 18 O H2O -dD H2O evolution after the gypsum saturation point (green dotted line) reflects the increasing importance of fractionation between free water molecules and water molecules within the hydration sphere of dissolved ions[START_REF] Sofer | Isotopic composition of hydration water in gypsum[END_REF] Gat, 1972, 1975). Note: MWL -Meteoric Water Line[START_REF] Craig | Deuterium and oxygen 18 variations in the ocean and the marine atmosphere[END_REF].

Fig. 4 .

 4 Fig. 4. Modelled co-evolution of: (a) salinity and d 18 O H2O , and (b) salinity and gypsum saturation state (X gypsum ) of mixtures of Atlantic water and continental runoff subject to evaporation with different evaporative conditions (salt-works versus Mediterranean Sea) and runoff chemistry (ion-free runoff versus Ca 2+ -and SO 4 2À

Fig. 7 .

 7 Fig 7c

Fig. 8 .

 8 Fig.8. Scenario proposed in this work for the formation of low-salinity gypsum as a side-product of the biogeochemical sulfur cycle.

Fig. 1 .

 1 Fig. 1. (a) Consensus chronostratigraphic model for the MSC events (Roveri et al., 2014a). Stage 3, here of interest, spans between 5.55 Ma and 5.332 Ma, the astronomical ages of the base of the Upper Gypsum Unit (following Manzi et al., 2009) and Trubi Formation (Van Couvering et al., 2000) in the Sicilian Eraclea Minoa section, respectively. CdB: Calcare di Base; PLG: Primary Lower Gypsum; RLG: Resedimented Lower Gypsum; UG: Upper Gypsum. (b), (c) Map of the Mediterranean region showing the two extreme and mutually exclusive paleoenvironmental scenarios proposed to have featured the Mediterranean during Stage 3 (see discussion in Chapter 7; modified after Krijgsman et al., 2018).

  Fig. 1. (a) Consensus chronostratigraphic model for the MSC events (Roveri et al., 2014a). Stage 3, here of interest, spans between 5.55 Ma and 5.332 Ma, the astronomical ages of the base of the Upper Gypsum Unit (following Manzi et al., 2009) and Trubi Formation (Van Couvering et al., 2000) in the Sicilian Eraclea Minoa section, respectively. CdB: Calcare di Base; PLG: Primary Lower Gypsum; RLG: Resedimented Lower Gypsum; UG: Upper Gypsum. (b), (c) Map of the Mediterranean region showing the two extreme and mutually exclusive paleoenvironmental scenarios proposed to have featured the Mediterranean during Stage 3 (see discussion in Chapter 7; modified after Krijgsman et al., 2018).

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Fig. 2 .

 2 Fig. 2. Map of the Mediterranean Basin (modified from Lofi, 2018) showing: a) the location of the key intermediate and deep basins as well as physical thresholds that influenced the connectivity history of the Mediterranean; b) the onshore (i.e. basins and/or sections) and offshore (DSDP/ODP/Industrial drill sites) localities where deposits attributed to MSC Stage 3 have been studied. Mixed assemblages of Paratethyan-like ostracods and foraminifera are known from all mentioned onshore localities and some offshore locations (see text). The present-day spatial extent of the MSC seismic units, except for the Lower Unit, is also shown. The paleogeography of the (Eastern and Central) Paratethys and of the North Aegean domain is contoured after Van Baak et al. (2017) and Krijgsman et al. (2020a), respectively. W-E onshore localities: 1-6 Betic Cordillera (SE Spain): 1-Marbella and 2-Malaga basins (Guerra-Merchán et al., 2010); 3-Sorbas Basin[START_REF] Roveri | The Terminal Carbonate Complex: the record of sea-level changes during the Messinian salinity crisis[END_REF](Roveri et al., , 2019a)); 4-Nijar Basin[START_REF] Fortuin | The Messinian of the Nijar basin (SE Spain): sedimentation, depositional environments and paleogeographic evolution[END_REF]; 5-Vera Basin[START_REF] Fortuin | The enigmatic Messinian -Pliocene section of Cuevas del Almanzora (Vera Basin, SE Spain) revisited-erosional features and strontium isotope ages[END_REF]; 6-Bajo Segura Basin[START_REF] Soria | The stratigraphic record of the Messinian salinity crisis in the northern margin of the Bajo Segura basin (SE Spain)[END_REF](Soria et al., , 2008a(Soria et al., , 2008b)); 7-Mallorca (Mas and Fornós, 2020); 8-Melilla Basin[START_REF] Rouchy | Late Messinian to Early Pliocene paleoenvironmental changes in the Melilla Basin (NE Morocco) and their relation to Mediterranean evolution[END_REF]; 9-Boudinar Basin[START_REF] Merzeraud | Sedimentology and sequence stratigraphy of the late-Messinian-Early pliocene continental to marine deposits of the Boudinar basin (North Morocco)[END_REF]; 10-Chelif Basin[START_REF] Rouchy | The end of the Messinian salinity crisis: evidences from the Chelif Basin (Algeria)[END_REF]; 11-Sahel area[START_REF] Frigui | Evidences of "Lago-Mare" episode around the Messinian-Pliocene boundary in eastern Tunisia (central Mediterranean)[END_REF]; 12-Aléria Basin and 13-Rhône Valley[START_REF] Carbonnel | La zone à Loxoconcha djaffarovi Schneider (ostracodes, Miocène supérieur) ou le Messinien de la vallée du Rhône[END_REF] 14-Piedmont Basin (Dela Pierre et al., 2011, 2016); 15-Po Plain[START_REF] Ghielmi | Sedimentary and tectonic evolution in the eastern Po-Plain and northern Adriatic Sea area from Messinian to Middle Pleistocene (Italy)[END_REF][START_REF] Ghielmi | Late Miocene-Middle Pleistocene sequences in the Po Plain-Northern Adriatic Sea (Italy): the stratigraphic record of modification phases affecting a complex foreland basin[END_REF][START_REF] Amadori | Restored topography of the Po Plain-Northern Adriatic region during the Messinian base level drop-Implications for the physiography and compartmentalization of the paleo-Mediterranean basin[END_REF]; 16-Fine Basin (Cava Serredi section;Carnevale et al., 2006a[START_REF] Carnevale | Did the Mediterranean marine reflooding precede the Miocene/Pliocene boundary? Paleontological and geochemical evidence from upper Messinian sequences of Tuscany, Italy. Paleogeogr[END_REF]. 17-21 Apennine system: Romagna sections (17,[START_REF] Roveri | Stratigraphy of the Messinian post-evaporitic stage in eastern-Romagna (northern Apennines, Italy)[END_REF], Trave section (18,[START_REF] Iaccarino | The Trave section (Monte dei Corvi, Ancona, central Italy): an integrated paleontological study of the Messinian deposits[END_REF], Maccarone section (19,[START_REF] Bertini | The Northern Apennines palynological record as a contribute for the reconstruction of the Messinian paleoenvironments[END_REF], Grossi et al., 2008;[START_REF] Sampalmieri | Paleoenvironments of the Mediterranean Basin at the Messinian hypersaline/ hyposaline transition: evidence from natural radioactivity and microfacies of postevaporitic successions of the Adriatic sub-basin[END_REF] Pellen et al., 2017), Colle di Votta (20)-Fonte dei Pulcini (21)-Stingeti (22) sections[START_REF] Cosentino | High-frequency cyclicity in the latest Messinian Adriatic foreland basin: insight into paleoclimate and paleoenvironments of the Mediterranean Lago-Mare episode[END_REF][START_REF] Cosentino | Orbitally forced paleoenvironmental and paleoclimate changes in the late postevaporitic Messinian of the central Mediterranean Basin[END_REF][START_REF] Cosentino | Refining the Mediterranean "Messinian gap" with high-precision U-Pb zircon geochronology, central and northern Italy[END_REF](Cosentino et al., , 2018)),Mondragone 1 well (23, Cosentino et al., 2006), Crotone Basin(24, Roveri et al., 2008a); 25-27 Sicily: Villafranca Tirrena (25) and Licodia Eubea (26) sections[START_REF] Sciuto | New reports of Paratethyan ostracods affinity from the Mediterranean Basin (Sicily, Italy)[END_REF]), Caltanissetta Basin (27, Manzi et al., 2009); 28-Corfu[START_REF] Pierre | Reconstruction of the paleoenvironmental changes around the Miocene-Pliocene boundary along a West-East transect across the Mediterranean[END_REF]; 29-Zakinthos(Karakitsios et al., 2017b); 30-Crete[START_REF] Cosentino | The late Messinian Lago-Mare episode in the Mediterranean Basin: preliminary report on the occurrence of Paratethyan ostracod fauna from central Crete (Greece)[END_REF]; 31-Cyprus[START_REF] Rouchy | Paleoenvironmental changes at the Messinian-Pliocene boundary in the eastern Mediterranean (southern Cyprus basins): significance of the Messinian Lago-Mare[END_REF] Manzi et al., 2016a); 32-Adana Basin[START_REF] Radeff | Stratigraphic architecture of the upper Messinian deposits of the Adana Basin (southern Turkey): implications for the Messinian salinity crisis and the Taurus petroleum system[END_REF].

  , left log) argued that every precessional/insolation peak must have an expression in the rock record.[START_REF] Manzi | A new facies model for the Upper Gypsum of Sicily (Italy): chronological and paleoenvironmental constraits for the Messinian salinity crisis in the Mediterranean[END_REF] agreed with the solution of Van

Fig. 3 .

 3 Fig. 3. (a), (b) Available astronomical tunings to astronomic curves of climatic precession (P), 100 kyr eccentricity (E) and 65 • N insolation curve (I) of Laskar et al. (2004) of the lithological cyclicity of onshore Stage 3sections (a) and of the seismic cycles and/or well logs (gamma ray and resistivity) of the MU in the Levant Basin (b). Tunings of onshore sections in (a) are carried out downward from the M/P boundary (conformable in all sections). Astronomically-tuned glacial (even numbers) and interglacial (odd numbers) stages (i.e. TG) as defined by Hodell et al. (1994) are also indicated.

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Fig. 4 .

 4 Fig. 4. Schematic overview of different chronostratigraphic models for some of the Messinian successions presented in Chapter 3. Note the large controversies in timing, duration and chronostratigraphic position of the main erosion phase between models in (a) and (b). Models in (a) follow the recently established MSC chronostratigraphic model of Roveri et al. (2014a), according to which the Mediterranean base-level dropped and halite deposited on sea floor during Stage 2 and the Upper Gypsum/Upper Evaporites-Lago-Mare sequence followed. Models in (b) were proposed following the alternative scenario of[START_REF] Andreetto | Alternate interpretation of the Messinian salinity crisis, controversy resolved?[END_REF][START_REF] Clauzon | Influence of Mediterranean sea-level changes on the Dacic Basin (Eastern Paratethys) during the late Neogene: the Mediterranean Lago Mare facies deciphered[END_REF], which envisaged two Lago-Mare episodes (LM1 and LM3) that occurred before and after the main Mediterranean drawdown event, during which LM2 was deposited in the deep desiccated basins[START_REF] Couto | Lago Mare and the Messinian salinity crisis: evidences from the Alboran Sea (S. Spain)[END_REF][START_REF] Popescu | Lago Mare episodes around the Messinian-Zanclean boundary in the deep southwestern Mediterranean[END_REF] see Roveri et al., 2008c and[START_REF] Andreetto | Migration of the dinoflagellate Galeacysta etrusca and its implications for the Messinian Salinity Crisis[END_REF] for further explanations). Note, in (b), the shifting of the position of the main erosional phase in Sicily through time as well as the time of the marine replenishment in the Apennines.

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Fig. 5 .

 5 Fig. 5. Sedimentary expression of Stage 3 from selected onshore Mediterranean localities. (a) Photograph from the Sorbas Basin showing the red continental sediments of the Zorreras member with intercalated white limestones (white arrows; from Andreetto et al., 2021). The conformable resting of the Zorreras Mb. above the Sorbas Mb. and underneath the Gochar Fm. of Pliocene age is also appreciable. Car for scale. (b) One typical lithological (and precessional) cycle of the Upper Mb. of the Feos Fm. in the Nijar Basin, here constituted by an ostracod-bearing, white and laminated mudstone bed overlain by an azoic fluvial sandstone (courtesy of Anne Fortuin). (c) Panoramic view of the Cuevas del Almanzora section (from Andreetto et al., 2021). Red rectangle indicates the position of the section straddling the Messinian (M)/Zanclean (Z) transition and studied by Fortuin et al. (1995), Stoica et al. (2016), Caruso et al. (2020) and Andreetto et al. (2021). Buildings for scale. (d) The sub-unit a of the Piedmont Basin composed of azoic grey mudstones grading into yellowish, mammal-rich overbank deposits. (e) WNW-ESE seismic profile in the Po Plain showing incised valleys filled during Stage 3 by suggested clastic deposits and sealed by deep-water turbidites in the Zanclean (modified from Amadori et al., 2018). (f) Typical aspect of the di Tetto/San Donato Formation in the Northern Apennines composed by grey mudstones (detail in the inset) with interbedded sandstone bodies (white arrows). The picture is taken from the Maccarone section. (g) The di Tetto Fm.-Colombacci Fm. transition in the Cusercoli area (Eastern Romagna, Fig. 2b), defined by the facies change underlined by the appearance of a fluvial conglomerate. This lithostratigraphic boundary also corresponds to substage 3.1/3.2 boundary of Roveri et al. (2014a). (h), (i), (l) Lithological cycles of the Upper Gypsum Unit in Eraclea Minoa (h), Siculiana Marina (i) and Polemi (l) sections. Cycles are several m-thick and primarily composed by beds of primary gypsum alternating with mudstones bearing Paratethyan ostracods (at least in Eraclea Minoa).

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Fig. 6 .

 6 Fig. 6. Photographs of the Miocene/Pliocene boundary (yellow lines) from selected onshore Mediterranean localities. (a) Erosive M/P transition in the Mendelín section (Malaga Basin). Note the conglomeratic lag draping the erosional surface and sharply overlain by foraminifera-rich marls. (b) Conformable stratigraphic contact between the uppermost Messinian sandstone of the Feos Fm. and the Zanclean biocalcarenites of the Cuevas Fm. in the Barranco de los Castellones section, Nijar Basin (hammer for scale; modified from Andreetto et al., 2021). (c) The Messinian/Zanclean boundary in the Pollenzo section (Piedmont Basin) marked by a characteristic black layer interbedded between Paratethyan ostracods-rich mudstones and marine foraminifera-rich marls (modified from Dela Pierre et al., 2016). (d) Uppermost segment of the Eraclea Minoa section (Caltanissetta Basin, Sicily) displaying the (non erosive) contact between the Pliocene Trubi Formation above and the sandy Arenazzolo Formation below. The inset is a close view of the transition, which occurs above a ~50 cm-thick burrowed mudstone horizon rich in Paratethyan ostracods and marine foraminifera. (e) Lago-Mare sediments in the Kalamaki section (Zakynthos) unconformable, through an erosional surface (i.e. the Messinian Erosional Surface, MES), over the PLG unit and also unconformable beneath the Trubi Fm. (modified from Karakitsios et al., 2017b). (f) Close view of the M/P boundary in the Pissouri Basin, where the foraminifera-rich Trubi marls lie above a black layer (paleosol according to Rouchy et al., 2001).

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1 )

 1 Stage 3 sedimentation follows a period of intensive tectonic and/or eustatic-driven erosion of the margins, as demonstrated by the frequent presence of erosional unconformities and/or chaotic Stage deposits (RLG unit); 2) Stage 3 lithologies are mostly terrigenous (conglomerates, sandstones and mudstones) and deposited in a variety of continental (fan delta, fluvial channels and alluvial plains) and shallow water environments (endorheic lakes or water bodies connected with the Mediterranean water mass is the riddle). Carbonate intercalations are sometimes present (e.g. Sorbas Basin and Colombacci limestones in the Apennines). Stage 3 gypsum is only found in deeper-water intermediate basins of Caltanissetta in Sicily and Cyprus. 3) A diversified fossil assemblage with Paratethyan affiliation (ostracods, dinocysts, mollusks) is commonly found in the shallow-water F. Andreetto et al.

Fig. 7 .

 7 Fig. 7. Seismic profiles from intermediatedeep Western Mediterranean basins containing MSC markers/units. (a) Seismic reflection line CAB01-104 from the WAB (modified from Booth-Rea et al., 2018). The line shows the variable geometry of the inferred M/P boundary, erosive in proximity of mud diapirs, (para)conformable in tectonically undisturbed sectors. Chaotic reflectors are occasionally imaged below the inferred M/P boundary. (b) Seismic profile SF12-09 imaging the lower slope of the south Algero-Balearic margin and part of the Algero-Balearic abyssal plain (modified from Mocnik et al., 2014). Here a high reflecting and horizontally stratified UU covers a thin layer of MU evidenced by salt diapirism. Note the concordant deformation of the UU and MU. (c) Line MS-39 from the abyssal plain of the Liguro-Provençal Basin showing the Messinian trilogy and non-erosive bottom and top surfaces (BS and TS; Dal Cin et al., 2016). Halokinesis of MU gives rise to domes that deform the UU and PQ units. (d) Interpreted seismic profile from the lowermiddle slope of the west Sardinian margin (modified from Dal Cin et al., 2016). Thin MU and UU are present on the lower slope, while on the middle slope (and upper slope here not shown) they converge in the margin erosion surface MES. (e) Line drawing of seismic line imaging from the Catalan margin (or Ebro Basin) to the abyss of the Liguro-Provençal basin (modified fromMaillard et al., 2011b). Note the pinch out of the MU in the Valencia Basin and of the UU in the Ebro Basin, which is MSC free. (f) Interpreted seismic profile Simbad crossing the depocenter of the CMD showing all the MSC units and erosional surfaces (modified fromRaad et al., 2021). (g) Interpreted seismic profile MYS40 illustrating the MU-UU-PQ units in the East-Sardinia Basin and Cornaglia Terrace, separated by the MSC deposits-free Quirra Seamounts (modified from[START_REF] Lymer | The Western Tyrrhenian Sea revisited: New evidence for a rifted basin during the Messinian Salinity Crisis[END_REF].

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Fig. 8 .

 8 Fig.8. Seismic profiles from intermediate-deep Eastern Mediterranean basins containing MSC markers/units (see Fig.7for abbreviations). (a) High-resolution seismic line MS27 imaging the PQ and the uppermost MSC's UU and MU in the Ionian Abyssal Plain and Gulf of Sirt (modified from[START_REF] Camerlenghi | Seismic markers of the Messinian salinity crisis in the deep Ionian Basin[END_REF]. Note how the MSC units are thinner, more difficult to distinguish and more deformed in the Sirt Abyssal Plain than in the IAP. PQ, UU and MU all onlap the structural highs of the Medina Ridge and VHS-2 sill. (b) High-resolution imaging of the lower part of the Plio-Quaternary (PQc unit) and upper part of the Messinian (UU) in the IAP[START_REF] Camerlenghi | Seismic markers of the Messinian salinity crisis in the deep Ionian Basin[END_REF]. The MSC-PQ boundary is a highly irregular surface, describing apparent V-shaped incisions (symbol V) of controversial origin (see subsection 4.6.1 for insights). Note the coherent deformation of PQc with the underlying MSC sequence and the absence of fluvial facies within the incisions (Unit II is made of lower Pliocene dolomitic marls recovered in Site 374 drilled nearby the seismic line; see text). (c) Multichannel seismic reflection profile MEM-07-203 running approximately parallel to the Malta Escarpment and showing the relationship between Unit 2 of[START_REF] Micallef | Evidence of the Zanclean megaflood in the eastern Mediterranean Basin[END_REF] with the overlying and underlying Zanclean and Messinian sediments, respectively (modified from[START_REF] Spatola | A single-stage megaflood at the termination of the Messinian salinity crisis: geophysical and modelling evidence from the eastern Mediterranean Basin[END_REF]. (d) Uninterpreted (left) and interpreted (right) seismic profiles showing the cyclic and channelized nature of the uppermost Messinian observed in the offshore Sirt Basin (modified from[START_REF] Bowman | A comprehensive review of the MSC facies and their origins in the offshore Sirt Basin, Libya[END_REF]. (e) Interpreted 2D regional WNW-ESE seismic profile crossing the continental shelf and offshore Levant Basin and the Herodotus Abyssal Plain[START_REF] Jagger | Tectono-stratigraphic evolution of the SE Mediterranean passive margin, offshore Egypt and Libya[END_REF]. Note the lateral continuity of the Messinian MU. (f) Seismic profile from the Levant Basin showing the 6 sub-units distinguished inside the MU as well as its lower (N-reflector) and upper (M-reflector) boundaries (modified from[START_REF] Gvirtzman | Intense salt deformation in the Levant Basin in the middle of the Messinian salinity crisis[END_REF]. (g) High-resolution seismic reflection image with wireline logs from Aphrodite-2 well illustrating that M-reflector previously considered as top evaporitic sequence and M/P boundary here consists of a ~100-m-thick unit (i.e. Unit 7 of[START_REF] Gvirtzman | Intra-Messinian truncation surface in the Levant Basin explained by subaqueous dissolution[END_REF] in which different layers are distinguished (modified from[START_REF] Gvirtzman | Intra-Messinian truncation surface in the Levant Basin explained by subaqueous dissolution[END_REF]. (h) Interpreted and uninterpreted seismic profiles imaging the Mavqi'im and Afiq formations described in the canyons on the Levant continental margin (modified from Ben[START_REF] Ben Moshe | Estimating drawdown magnitudes of the Mediterranean Sea in the Levant basin during the Lago Mare stage of the Messinian Salinity Crisis[END_REF].

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Fig. 9 .

 9 Fig. 9. Photomicrographs of the key microand macrofossils featuring Stage 3 sediments. (a) Scanning electron microscope (SEM) photographs of the more common Paratethyan ostracod species in substage 3.1 and 3.2 sediments (compiled from Stoicaet al., 2016, Cosentino et al., 2018[START_REF] Sciuto | New reports of Paratethyan ostracods affinity from the Mediterranean Basin (Sicily, Italy)[END_REF]. (b) Photomicrographs of the Paratethyan dinoflagellate cyst Galeacysta etrusca under the optical microscope (left) and SEM (right) (modified from Do[START_REF] Couto | Lago Mare and the Messinian salinity crisis: evidences from the Alboran Sea (S. Spain)[END_REF][START_REF] Andreetto | Migration of the dinoflagellate Galeacysta etrusca and its implications for the Messinian Salinity Crisis[END_REF].

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Fig. 10 .

 10 Fig. 10. Isotopic record of MSC Stage 3 for the Mediterranean Basin. (a) Compilation of MSC Stage 387 Sr/ 86 Sr isotope data sourced from ostracod valves and gypsum crystals (see Supplementary material 1 and subsection 6.1 for references). Data are plotted with the global 87 Sr/ 86 Sr seawater curve(McArthur et al., 2012). Error bars indicate analytical error, which is so small in some cases that no error bars are visible at this scale. To not complicate the figure, horizontal error bars have not been added for the sections/cores unprovided of a chronostratigraphic framework and for which age uncertainties are present (i.e. all but Nijar and Vera basins, Eraclea Minoa and onshore Cyprus; see Fig.3). Note that none of the 87 Sr/ 86 Sr isotope ratios but one from Nijar plot on the ocean curve. In the inset is shown the Mediterranean Sr record for the entire MSC as well as the time-equivalent Eastern Paratethys record (modified afterAndreetto et al., 2021). (b) Plot of δ 34 S SO4 and δ 18 O SO4 in Stage 3 gypsum and anhydrite beds from onshore and offshore localities (see Supplementary material 1 and subsection 6.2 for references). No measures are available from the marginal basins, where gypsum did not deposit during Stage 3. The dark blue and black rectangles represent the sulfate isotopic composition of the Global Messinian ocean and Stage 1 (PLG) evaporites, respectively. The light blue area represents the sulfate isotopic composition of mixtures of Messinian marine waters with non-marine sources. The red area represents the isotopic composition of the residual sulfate ion in a basin where marine Messinian sulfate is consumed by microbial sulfate reduction to produce H 2 S. The arrow represents the isotope trajectory of dissolved sulfate resulting from the mixing of residual 34 Senriched sulfate produced by MSR and 34 S-depleted sulfate produced by H 2 S oxidation. All the published δ 34 S SO4 and δ 18 O SO4 values are provided corrected with the fractionation factors δ 34 S=+1.65‰ and δ 18 O=+3.5‰ to smooth the isotopic fractionation effects experienced by dissolved sulfate and to reason on values reproducing the isotopic composition at the time of gypsum precipitation. (c) δD isotopes of C 29 and C 31 n-alkanes and C 37 and C 38 long chain alkenones recorded in the Stage 3 gypsums and marls of the Eraclea Minoa section (modified fromVasiliev et al., 2017). Blue lines indicate the values recorded in the present day lacustrine settings for the n-alkanes[START_REF] Sachse | δD values of individual n-alkanes from terrestrial plants along a climatic gradient -implications for the sedimentary biomarker record[END_REF] and in the alkenones from the Mediterranean in the recent times[START_REF] Van Der Meer | Hydrogen isotopic compositions of long-chain alkenones record freshwater flooding of the Eastern Mediterranean at the onset of sapropel deposition[END_REF]. Error bars indicate standard errors of the mean.

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Fig. 11 .

 11 Fig. 11. (a), (b) Schematic W-E profiles across the Mediterranean Basin showing the contrasting paleoenvironmental, paleohydrological and paleoconnectivity interpretations proposed for Stage 3. When a water flow is present (green arrow) from and/or to an extra-Mediterranean water mass (i.e., A: Atlantic Ocean; I: Indian Ocean; P: Eastern Paratethys), the direction of the arrow gives the direction of flow. For simplicity, water added by the major and local rivers is not shown, but it adds to the hydrological budget at any time in each scenario. Note the main difference between the isolated (a) and density-stratified (b) scenario lies in the connectivity framework (Atlantic connection closed and negligible influence from the Paratethys in the isolated scenario; influence from both Atlantic and Paratethys in the density-stratified scenario), which affects the position of the base level of the Mediterranean water mass and its hydrochemistry (see extensive discussion in subsection 7.2). Abbreviations: Sp.: SE Spain; V-B: Valencia Basin; Tyr: Tyrrhenian Basin; Calt: Caltanissetta Basin; IAP: Ionian Abyssal Plain; GS: Gulf of Sirt; Cyp: Cyprus; Ada: Adana Basin. See Fig. 2 for the geographic position of each basin. (c) Schematic plot showing the evolution of the Mediterranean base-level during Stage 3 according to both the isolated (red line) and half-full (black line) scenarios. The critical sills for controlling intra-and extra-Mediterranean connectivity are also shown.

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  Crown Copyright © 2022 Published by Elsevier B.V. This is an open access article under the CC BY license (http://creativecommons.org/licenses/by/4.0/).

  Manzi et al. (2009)d e s c r i b e d three repeated facies associations of the Upper Gypsum at Eraclea Minoa: two distinct gypsumdominated facies associations (B

Fig. 1 .

 1 Fig. 1. Simplified geological map of Sicily (upper left figure; modified from Manzi et al., 2021) and NW-SE-oriented geological cross section (upper right figure; modified from Manzi et al., 2021) showing the geographical and stratigraphic position of the Eraclea Minoa section, respectively. Asterisks in the upper left figure indicate the position of other sections mentioned in the text (1: Balza Soletta, Maniscalco et al., 2019; 2: Montedoro, Rouchy and Caruso, 2006; 3: Siculiana Marina, Caruso and Rouchy, 2006). Below is a panoramic view of (part of) the Eraclea Minoa section. Four gypsum beds (III, IV, V and VI) overlooking the Eraclea Minoa village stand out in the foreground. The cliff of Cattolica Eraclea/Capo Bianco in the background is mainly composed of the white marly deposits of the Pliocene Trubi Fm. unconformably overlain by the Agrigento Fm. of Lower Pleistocene age. Locally, the topmost part of the UG and/or the Arenazzolo Fm. crop out at the cliff foot.

Fig. 2 .

 2 Fig. 2. Lithological log of the Eraclea Minoa section (modified from Manzi et al., 2009). An ostracod stratigraphic distribution chart (Fig. 2) was constructed integrating the results of Grossi et al. (2015)w i t h seventeen marl samples from the base of cycle 3 provided by Elsa Gliozzi and Domenico Cosentino and prepared at the Earth Science Department of the Universitá degli Studi Roma Tre with standard micropaleontological methods of washing, sieving and drying. New and published (Grossi et al., 2015; García-Veigas et al., 2018) 87 Sr/ 86 Sr ratios measured on gypsum crystals and ostracod valves and sulfur and oxygen (in sulfate) isotope profiles from García-Veigas et al. (2018)a r e also plotted. The bottom diagram is a close view of the 87 Sr/ 86 Sr isotopic composition of the gypsum beds I-VII. Data are plotted against the coeval oceanic ratio (after McArthur et al., 2012) and the composition of a system mixing river water from Nile and Rhône (0.707427; Topper et al., 2011) and the Eastern Paratethys (0.7084-0.7085 during the latest Messinian; Grothe et al., 2020). The lack of overlap between gypsum and water data are evidence that the gypsum-precipitating water was a combination of all three sources mentioned above. (1): This study; (2): Grossi et al. (2015); (3): García-Veigas et al. (2018).

Fig. 3 .Fig. 4 .

 34 Fig. 3. a: Schematic representation of a complete UG depositional cycle. Four sub-units (A to D) are distinguished alongside the three facies associations of Manzi et al. (2009). See legend in Fig. 7a. b: Outcrop view of cycle 6 which contains all four sub-units. Note the sharp boundaries (red dashed lines) between the sub-units. c: Close up of sub-units A to C of cycle 5. Yellow dashed lines indicate the sub-unit boundaries. d: Alternations of white diatomites (Di) and greenish marls (Ma) in sub-unit A of cycle 5. e: Close-up of sub-unit B in cycle 5 showing planar-laminated gypsarenite (Gya), greenish marls (Ma) and balatino gypsum beds (Bal). The laminae of the balatino gypsum locally show small scale folding. f-g: Transition from sub-unit B to C (f) and sub-unit C to D (g) from cycle 7 as exposed along the beach at the foot of the cliff Cattolica Eraclea/Capo Bianco cliff. Yellow arrows in Fig. 3g point to brownish litharenites interbedded with gray marls of sub-unit D. (For interpretation of the colors in the figure(s), the reader is referred to the web version of this article.)

Fig. 5 .

 5 Fig. 5. a: Mass balance equations and strontium parameter values employed to model the 87 Sr/ 86 Sr ratio of a water mass mixing Atlantic, Eastern Paratethys and large circum-Mediterranean rivers (Nile and Rhone) waters in various proportions. 87 Sr/ 86 Sr Atlantic refers to the average value of the range of 87 Sr/ 86 Sr ratios (0.709017-0.709027) measured on calcareous material assigned to the time interval 5.55 and 5.30 Ma (McArthur et al., 2012). 87 Sr/ 86 Sr Nile+Rhône , [Sr] Atlantic and [Sr] Nile+Rhône are from present-day measurements (e.g. Topper et al., 2011). 87 Sr/ 86 Sr Paratethys corresponds to the midpoint value of the range of 87 Sr/ 86 Sr ratios (0.7084-0.7085) derived from ostracods recovered from Stage 3-equivalent sediments in the Caspian Basin (Grothe et al., 2020). b-c: Ternary plots of the 87 Sr/ 86 Sr ratios modeled for a system mixing Atlantic seawater, riverine freshwater from Nile and Rhône and brackish water from the Eastern Paratethys using Two values of [Sr] Paratethys , which is the major unknown in the model, are employed. In b, [Sr] Paratethys is 0.296 mg/l, similar to the combination Nile+Rhône, whereas in c the [Sr] Paratethys is slightly higher (0.5 mg/l). The range of 87 Sr/ 86 Sr ratios measured for the Eraclea Minoa Upper Gypsum layers (pink area) is also plotted to facilitate the comparison with the model outputs. The ternary diagram is generated using the free and open-source software https://www.ternaryplot .com/.

Fig. 6 .

 6 Fig. 6. Snapshot of the Atlantic-Mediterranean-Eastern Paratethys region during times of Stage 3 gypsum precipitation (i.e. sub-units B and C) in Eraclea Minoa. The gypsumprecipitating Lago-Mare water mass is drawn following the present-day spatial distribution of the Upper Gypsum (UG) in the intermediate basins (Sicily and Cyprus) and Upper Unit (UU) in the deep basins (taken from Lofi, 2018). A different water mass was occupying the Stage 3 gypsum-free Northern Aegean, Adriatic and Western Alborán Basin (WAB), likely due to the presence of structural sills. The insets show a schematic close view of the Atlantic-Mediterranean and Mediterranean-Paratethys connections.River water was always draining into the Mediterranean with variable discharge at different moments during a precession/insolation cycle(Simon et al., 2017).

Fig. 7 .

 7 Fig.7. a: Lithofacies of a complete UG cycle as a function of precession cycle and Mediterranean base-level according to[START_REF] Manzi | A new facies model for the Upper Gypsum of Sicily (Italy): chronological and paleoenvironmental constraits for the Messinian salinity crisis in the Mediterranean[END_REF]a n d this study. A comparison with the PLG interpretation is also shown (re-drawn afterRoveri et al., 2008). The question marks in the figure to the left refer to the uncertainty over the paleodepth during deposition of sub-units A, B and C. Note that our new interpretation of the UG challenges the long-lasting idea that PLG and UG reflect similar depositional mechanisms and paleoenvironmental conditions[START_REF] Manzi | A new facies model for the Upper Gypsum of Sicily (Italy): chronological and paleoenvironmental constraits for the Messinian salinity crisis in the Mediterranean[END_REF] Roveri et al., 2014a). This is consistent in a number of ways: (1) UG and PLG sequences comprise different facies (UG contains more abundant balatino gypsum and is missing the banded and branching selenite lithofacies of the PLG). (2) Euxinic shales interbedded with the PLG contain possibly in-situ marine fossils (e.g.Corbí and Soria, 2016), whereas the marls interbedded with the UG contain brackish and shallow water ostracods and brackish water foraminifera such as A. tepida (Fig.2;Grossi et al., 2015). This is likely to reflect a greater proportion of marine water in the Mediterranean during Stage 1. (3) 87 Sr/ 86 Sr isotopes are consistent with the above interpretation, being, on average, close to, or within error of coeval ocean values on stage 1 gypsums (average of 0.708941;Lugli et al., 2010; Reghizzi et al., 2018) and substantially higher than those measured on Stage 3 gypsum samples (average of 0.708752; Fig.7a). (4) The Stage 3 gypsum 87 Sr/ 86 Sr data are either constant within the layer or show a subtle decline from base to top (Figs.5a, 7). This is in contrast with a larger base to top increase in 87 Sr/ 86 Sr ratio in Stage 1 gypsum layers from the Vena del Gesso Basin (Northern Apennines;Reghizzi et al., 2018; Fig. 7a). (5) The PLG is largely restricted to marginal basins (e.g.Lugli et al., 2010; García-Veigas et al., 2018), whereas the UG is only found in intermediate-deep basins (Fig.6). b: Summary outline of the main sedimentological (i), paleontological (ii) and geochemical (iii) data for each sub-units, their interpretation in terms of salinity (iv) and relationship with the base-level curve (v) and their paleoenvironmental and paleohydrological interpretation proposed in this work (vi).

11 Fig. 8 .

 118 Fig. 8. a: Revised astronomical tuning of the Upper Gypsum Unit at Eraclea Minoa according to the newly proposed phase relationships between the UG lithofacies and the astronomical forcing (namely 65 • N summer insolation and 100 kyr eccentricity curves). Starting from the conformable M/P boundary, we correlate the seven gypsum beds to successive transitions from insolation maxima to minima avoiding the insolation peaks associated with the 100-kyr eccentricity minima at ∼5.38 Ma and ∼5.50 Ma. Due to the low amplitude of insolation in these eccentricity low, these minor peaks are not expected to correspond to a gypsum bed in the sedimentary record (van der Laanet al., 2006). It is more likely that these low eccentricities corresponded to prolonged phases of very low base level in the Mediterranean (Fig.8a). Pursuing this approach, all gypsum beds follow well pronounced insolation maxima and the thicker, gypsum-free, Cyprideis-rich sub-units D between gypsums II-III and VI-VII correlate with an eccentricity minimum. The overall Mediterranean run-off budget fromSimon et al. (2017), with included (violet curve) and excluded (green curve) supply from the Chad lake, is also shown. b-e: Cartoons showing the proposed evolution of the Mediterranean environments and hydrology during one precession cycle of MSC Stage 3.

Fig. 1 .

 1 Fig. 1. PLG in the Mediterranean and in the VdG basin. A. Sulfur δ 34 S SO4 and δ 34 S IS isotopic curve of the PLG succession of the VdG basin, calibrated with the 100ka orbital eccentricity and precession curves (Laskar et al., 2011). Note the nearly invariant δ 34 S SO4 isotopic variations compared to δ 34 S IS . B. Panoramic view of the 220 m thick VdG succession constituted by an alternation of thick gypsum beds and marl intervals (PLG unit). C. Distribution of Messinian evaporites in the Mediterranean and the location of the VdG basin (pink star) (see supplementary section 2.c for references). (For interpretation of the colors in the figure, the reader is referred to the web version of this article.)

Fig. 2 .

 2 Fig. 2. Macro-and microscopic characteristics of VdG PLG gypsum deposits. A. The basal part of cycle PLG5, composed of a basal dark marly interval followed by a thin whitish carbonate layer, in turn overlain by a gypsum bed consisting of dm-sized vertically oriented selenite crystals (massive selenite lithofacies). B. Close-up of the banded selenite facies (cycle PLG7). C. Close-up of a vertically oriented selenite crystal (massive selenite lithofacies). Note the turbid portion in the re-entrant angle of the twin. D. Photomicrograph (transmitted light) of a twinned selenite crystal showing the lamination in the re-entrant angle of the twin. E. Photomicrograph (transmitted light) of a gypsum twin with curved filaments aligned to the vertical growth bands. Lamination in the re-entrant angle is also shown. F. Detail of (E) with filaments aligned along a vertical growth band. G, H, I. Photomicrographs in transmitted light (G and H) and UV light (I) showing the filaments inside gypsum. Pink arrows show tiny iron sulfide globules within the filaments. Note in (I) the high autofluorescence of the filaments. J, K, L. Photomicrographs in transmitted (J, K) and reflected (L) light of clay-rich aggregates in the turbid laminae of gypsum crystals. The aggregates are interpreted as marine snow floccules and contain abundant sulfide grains (arrows). (For interpretation of the colors in the figure, the reader is referred to the web version of this article.)

3. 1 .

 1 δ 34 S and δ 18 O of sulfate

Fig. 3 .

 3 Fig. 3. Stable sulfate isotopic composition of VdG gypsum.

Fig. 4 .

 4 Fig. 4. Multiple Sulfur isotopic composition of sulfate and sulfide minerals constituting PLG deposits.

Fig. 7 .

 7 Fig. 7. Schematic representation of the biogeochemical conditions that accompanied PLG deposition. (A) Deposition of organic rich marls during humid phases (precession minima). (B) Deposition of gypsum in the margins and marls in the depocenter during arid phases (precession maxima). The relative size of the arrows represents the approximate magnitude of fluxes. (C) Details of a gypsum twin. (For interpretation of the colors in the figure, the reader is referred to the web version of this article.)

  highly debated (seeAndreetto et al., 2021 for a detailed review). Regardless of this, similarly to more marginal deposits, the strontium isotopic composition of the ostracods found in deep sea drill cores are all well below global Messinian ocean strontium isotopic composition, suggesting that large amount of non-oceanic strontium fed the deep basins of the Mediterranean at the very end of MSC phase 3 (Figure2. 22).

	Chapter 2-Material and Methods
	___________________________________________________________________________________________
	or reworked is
	76

Chapter 2-Material and Methods ___________________________________________________________________________________________ 77 2.2.2 Sampled ODP and DSDP sites Figure 2. 31 Lithological log of the cores studied in this work.

  this work we analysed a total of 45 samples from DSDP and IODP drilling cores (see black dots in Figure 2. 31). Samples were chosen from the different representative lithofacies of each

	Chapter 2-Material and Methods Chapter 2-Material and Methods
	___________________________________________________________________________________________ ___________________________________________________________________________________________
	Isotope Geochemistry
	Chemical extraction, measurements and interpretation
	___________________________________________________________________________
	2.3 Sulfur and oxygen isotopes
	2.3.1 Chemistry and isotopic measurements of multiple sulfur isotopes
	2.3.1.1 Sulfur extraction
	Reductive distillation reactions allowed the extraction of the sulfur phases from our samples
	84 85

studied core. In gypsum, we analyzed the multiple sulfur and oxygen isotope composition of the sulfate ion (δ 34 SSO4, Δ 33 SSO4 and δ 18 OSO4, respectively). The multiple sulfur isotope composition of iron sulfide δ 34 SIS and Δ 33 SIS) was also analyzed in both gypsum and marly interbeds. Sulfide was chemically extracted from the gypsum layers (of which they are very minor component) and the marls (see 2.3 for the full procedure). In total, we analyzed: a) 29 samples of gypsum from site 134, 654, 374 and 376 for the oxygen isotopic composition of sulfate (4 anhydrite, 12 laminated gypsum, 7 clastic gypsum, and 4 porfiroblastic gypsum). b) 29 samples of gypsum for the multiple sulfur isotopic composition of sulfate (4 anhydrite, 13 laminated gypsum, 7 clastic gypsum and 4 porfiroblastic gypsum). c) 42 sulfide extracts (5 anhydrite, 11 laminated gypsum 8 clastic gypsum, 4 porfiroblastic gypsarenite and 13 interbedded marls).

and their transformation into hydrogen sulfide H2S and then silver sulfide Ag2S. Calcium sulfate minerals (gypsum or anhydrite) and marls samples were first crushed into 100µm

  and for other studies measuring ∆ 33 S (Au Yang et al., 2016 and references therein). Given the large incertitude associated to ∆ 36 S measurements during t he course of our study (see below), we don't further consider it in our interpretations.

	Chapter 2-Material and Methods
	___________________________________________________________________________________________
	The error bar on the whole procedure for sulfate and sulfide extraction and measurements
	were respectively calculated from replicate extraction and analyses of barite (n=7), gypsum
	(n=6) and marls (n=9) internal standards. We obtained 0.2 and 0.004‰ of error on δ 34 S and
	Δ 33 S of sulfate and 0.4 and 0.008‰ for δ 34 S and Δ 33 S respectively for sulfides.
	2.3.2 Chemistry and isotopic measurements of oxygen isotopes of sulfate
	Oxygen isotope measurement of the sulfate molecule of gypsum or anhydrite was obtained
	by analyzing barite BaSO4. BaSO4 was produced by re-precipitating the SO4 2-ion of the calcium
	93

IAEA-S1 VS CDT (n=29, 1σ), δ 34 S= -0.3 ±0.1‰, ∆ 33 S= 0.084±0.003‰, ∆ 36 S= -0.705 ±0.542‰ IAEA-S2 VS CDT (n=15, 1σ), δ 34 S= 22.4 ±0.1‰, ∆ 33 S= 0.032±0.006‰, ∆ 36 S= -0.236 ±0.256‰ IAEA-S3 VS CDT (n=13, 1σ), δ 34 S= -32.4 ±0.1‰, ∆ 33 S= 0.084±0.004‰, ∆ 36 S= -0.922 ±0.391‰ sulfate mineral in a BaCl2 solution in the Stable isotope laboratory in IPGP. The isotopic analyses were conducted at two different places. (1) Oxygen isotopic composition of the sulfate ions of the Vena del Gesso gypsum was analysed at University of Burgundy on a Elementar vario PYRO cube elemental analyzer in-line with an Isoprime 100 mass spectrometer in continuous flow mode; the analytical error is ±0.4‰ (2σ) based on replicate analyses of the international barite standard NBS-127. DSDP and IODP gypsum and anhydrite samples were analyzed at the LOCEAN laboratory, Université de Paris on a ThermoScientific Delat V Advantage Isotope Ratio Monitoring Mass Spectrometer (IRMS) via pyrolysis as explained in

  Then we calculated the derivative of λ as a function of δ 34 S and ∆ 33 S as follow:

														Chapter 2-Material and Methods
	___________________________________________________________________________________________
	𝜕𝜆 #$% 𝜕𝜀 6; 𝑆
	= -	ln B1 + ∆ 66 𝑆 8(+9+:; + 1000 E-1 + 31 + 1000 31 + 𝜀 6; 𝑆 1000 < × ln 31 + 1000 𝜀 6; 𝑆 1000 𝜀 6; 𝑆 < < ( " KLM	H /1000J
	+	1000	⎝ ⎜ ⎛	1 +	31 + ∆ 66 𝑆 8(+9+:; + 1000 E-1 + 31 + 𝜀 6; 𝑆 1000 < 9OP" KLM × 0.515 𝜀 6; 𝑆 1000 < 1000	" KLM	H	⎠ ⎟ ⎞	× ln 31 +	𝜀 6; 𝑆 1000 <
	and
	𝜕𝜆 #$% 𝜕∆ 66 𝑆	=	1000	⎝ ⎜ ⎛	1 +	1 ∆ 66 𝑆 8(+9+:; + 1000 E-1 + 31 + 1000	𝜀 6; 𝑆 1000	<	" KLM	H	⎠ ⎟ ⎞	× ln 31 +	𝜀 6; 𝑆 1000 <
														95

Table 2

 2

. 1 Sulfur fractionation factors used in the numerical model

rate Fgyp that extracts sulfate from the marginal basin:

			Chapter 2-Material and Methods
		___________________________________________________________________________________________
	I.	gypsum accumulation with SO4 2-+ Ca 2+ + 2H2O → Ca.SO4.2H2O	(1)
	II.	dissimilatory sulfate reduction into sulfide (MSR), carried out by sulfate reducing
		microbes, with rate Fmsr:
		SO4 2-+ 2CH2O → H2S + 2HCO3 -	(2)
	III.	oxidation of
			99

sulfide produced by MSR to sulfate, with rate Fox, which can take place:

	(iii.1) chemically, using molecular oxygen or metal oxides as electron acceptors
	H2S + 2O2 → SO4 2-+ 2H +	(3)
	Or
	H2S + 3/2O2 → SO3 2-+ 2H +	(3.1)
	and
	SO3 2-+ 1/2O2 → SO4 2-(Avrahami and Golding, 1968; Zhang and Millero, 1993)	(3.2)
	Or
	4FeOOH + 1/2HS -+ 8H + → 4Fe 2+ + 1/2SO4 2-+6H2O (Seitaj et al., 2015)	(4)
	Or
	4MnO2 + HS -+ 7H + → 4Mn 2+ + SO4 2-+ 4H2O (Aller and Rude, 1988)

  0

Table 2 . 1

 21 Sulfur fractionation factors used in the numerical model *see next paragraph Although Eldridge et al. (2021) have calculated fractionation factors for both 33 S and 34 S isotopes and for the different species of solid sulfide possibly generated

	Chapter 2-Material and Methods
	___________________________________________________________________________________________

Eldridge et al., 2021) (10.1)

fixed values (see Table

2

. 1), except for MSR for which isotope fractionation is a function of the cellular rate of microbial sulfate reduction (see section 2.4.2.2). In Table

2

. 1, Rx are the isotope ratios ( x S/ 32 S, where x = 34 or 33) of the sulfur fluxes involved in the fractionation processes: Roc is the sulfur isotope ratio of the sulfate input from the ocean, Rriv is the sulfur isotope ratio of sulfate input from the rivers, Rmb is the sulfur isotope ratio of sulfate in the marginal basin, Rgyp is the sulfur isotope ratio of gypsum precipitated in the marginal basin, Rmsr is the sulfur isotope ratio of H2S produced by MSR, Rox is the sulfur isotope ratio of S° produced by oxidation of H2S, Rdo is the sulfur isotope ratio of SO4 2-produced by disproportionation of S° (or SO3 2-) and Rdh is the sulfur isotope ratio of H2S produced by disproportionation of S°. Chapter 2-Material and Methods ___________________________________________________________________________________________ 102 ___________________________________________________________________________________________ 103 Solid Sulfide Burial :

  𝐹 á# + 𝐹 ~+ 𝐹 âáå = 𝐹 çéè + 𝐹 àÉâ + 𝐹 ~ê% (18) and 𝐹 á# × 𝑅 ~Ç + 𝐹 ~× 𝑅 ~+ 𝐹 âáå × 𝑅 âáå = 𝐹 çéè × 𝑅 çéè + 𝐹 àÉâ × 𝑅 àÉâ + 𝐹 ~ê% × 𝑅 àë

	Chapter 2-Material and Methods

___________________________________________________________________________________________ 106

For the SO4 2-box, the two mass balances are:

  𝑅 àÉâ + 𝐹 ñ × 𝑅 ñ = 𝐹 èé × 𝑅 èé + 𝐹 ~Ü × 𝑅 ~Ü

		Chapter 2-Material and Methods
	___________________________________________________________________________________________
	𝐹 àÉâ + 𝐹 ñ = 𝐹 èé + 𝐹 ~Ü and	(24)
	𝐹 àÉâ ×
		107

Modelled isotope composition of VdG deposits compared to measured values. Data

	Chapter 3-Cryptic sulfur cycling during the formation of giant gypsum deposits
	_____________________________________________________________________________________________________
	(i) Figure 3. 7 points MSR only (without reoxidation of sulfide).
	in (A) represent the Δ 33 S vs δ 34 S of sulfate and sulfide of gypsum (massive, banded, and branching
	(ii) selenite) and marl samples from the PLG unit of the VdG Basin. Data points in (B) represent the isotopic MSR coupled to direct abiotic oxidation of sulfide to sulfate.
	fractionation calculated for paired sulfate and sulfide in the various gypsum lithofacies (massive,
	(iii) banded, and branching selenite). Individual curves represent the modelled isotopic compositions in (A) MSR coupled to phototrophic oxidation of sulfide to sulfate or elemental sulfur.
	and net fractionations factors in (B) for different values of fox (solid lines) and ƐMSR (dashed lines) the
	(iv) extent of isotopic fractionation during MSR. ƐMSR approaches -70‰ when thermodynamic equilibrium MSR coupled to partial oxidation of sulfide to elemental sulfur by
	is reached and fox=1, indicating that all the sulfide produced by MSR is reoxidized. Note that in both chemolithotrophs and subsequent disproportionation of elemental sulfur. diagrams most of the data fall in the area defined by fox>0.85, indicating the occurrence of nearly
	(v) complete reoxidation of sulfide, and supports the occurrence of ongoing cryptic sulfur cycling during MSR coupled to partial oxidation of sulfide to sulfite and subsequent gypsum formation. Most of the data fall in the area defined by ƐMSR <-50‰, indicating isotopic
	disproportionation of sulfite. fractionation that deviated from thermodynamic equilibrium. (Error bars are 1σ).
	Supplementary section 3.8.2 provides full details on the modelling and results of the various
	biogeochemical scenarios.In all scenarios (i) to (v), the entire range (from 0 to 1) of fox was
	explored (fox refers to the fraction of H2S produced by MSR being reoxidized; see section 2.4
	for a full description of the model equations and parameters). The isotope fractionation
	factors are fixed for all processes except for MSR, for which the full ƐMSR and λMSR ranges have
	been explored (see section 2.4.2.2 for further explanation).
	3.5.3.2 Cryptic sulfur cycling during the formation of VdG gypsum layers
	The best fit of the model results to the analytical dataset (
	Figure 3. 7) is obtained with fmsr = 0.8 (representing the fraction of sulfur entering the system
	as SO4 and subsequently reduced by MSR, see Supplementary section 3.8.2.1), and scenario
	(iv), i.e. the partial oxidation of sulfide to elemental sulfur by chemolithotrophs and
	subsequent disproportionation of elemental sulfur (Figure 3. 8C and
	Figure 3. 7) (See Supplementary section 3.8.2.3 for further discussion on the reoxidative
	processes involved in sulfur cycling).
	138

Table 3 .

 3 1 Concentrations and multiple sulfur isotopic composition of Iron sulfide from the Vena del Gesso gypsum and marly interbeds. Error is 1σ.

PLG Cycle Facies Sample name Δ 33 SIS ±0.008‰ δ 34 SIS ±0.4‰ [IS] ±100ppm

			Chapter 3-Cryptic sulfur cycling during the formation of giant gypsum deposits Chapter 3-Cryptic sulfur cycling during the formation of giant gypsum deposits
	_____________________________________________________________________________________________________ _____________________________________________________________________________________________________
	7 12	Laminated gypsum Branching selenite	La VII-02 Br XII-06	0.120 0.128	-36.35 -34.80	234
	7 13	Massive selenite Banded selenite	Ma VII-03 Ba XIII-03	0.068 0.127	-33.08 -37.12	502
	7 13	Banded selenite Branching selenite	Ba VII-04 Br XIII-04	0.034 0.016	-9.26 -12.84	35 166
	7 14	Branching selenite Massive selenite	Br VII-05 Ma XIV-02	0.054 0.036	-17.83 6.06	97 117
	7 14	Branching selenite Banded selenite	Ba VII-06 Ma XIV-03	0.126 0.023	-27.16 -4.98	219
	1 8 8 14 15 2 8 15 2 8 15 3 3 9 9 15 16 4 9 16 4 4 5 5 5 6 6 6 6 7 12 12 12 11 11 11 11 10 10 10 16 10 16	Marl Massive selenite Branching selenite Marl Massive selenite Banded selenite Massive selenite Massive selenite Branching selenite Banded selenite Massive selenite Marl Marl Massive selenite Branching selenite Marl Massive selenite Marl Banded selenite Massive selenite Massive selenite Banded selenite Marl Massive selenite Marl Massive Banded selenite Marl selenite Banded selenite Massive Branching Marl selenite selenite Branching selenite selenite Banded selenite Massive Banded Marl selenite selenite Branching Marl Massive selenite Massive selenite Banded selenite Banded selenite	Pe VIII-01 Ma VIII-02 Br XIV-04 Pe XV-01 Ma I-01 Ba VIII-03 Ma XV-02 Ma II-02 Br VIII-04 Ba XV-03 Ma II-07 Pe III-01 Pe IX-01 Ma IX-02 Br XV-04 Pe XVI-01 Ma III-01 Pe IV-01 Ba IX-03 Ma XVI-02 Ma IV-02 Ba IV-05 Pe V-01 Ba V-06 Pe VI-01 Br VI-04 Pe VII-01 Ba XII-05 Ma XII-04 Pe XII-03 Ma VI-03 Br XI-04 Ma VI-02 BA XI-03 Ma XI-02 Pe XI-01 Ma V-03 Br X-04 Pe X-01 Ma X-02 Ma XVI-02 bis Ba X-03 Ba XVI-05	0.125 0.040 0.061 0.121 0.045 0.086 0.029 -0.017 0.089 -0.008 0.025 0.143 0.130 0.061 0.100 0.014 0.046 0.132 0.122 0.031 0.025 0.120 0.142 0.071 0.161 0.034 0.119 0.127 0.026 0.133 0.082 0.099 0.037 0.149 0.021 0.143 0.018 0.074 0.132 0.045 0.009 0.047 0.027	-32.08 -29.46 -26.00 -36.86 -3.63 -27.46 -13.96 -18.06 -25.71 -14.78 0.36 -33.87 -30.26 -36.01 -21.34 -17.84 8.88 -35.37 -30.47 -2.73 0.34 -32.56 -33.29 -27.89 -32.54 -12.74 -31.12 -32.09 -1.11 -34.31 -21.86 -22.51 -13.90 -33.83 -0.26 -32.00 -11.16 -18.19 -31.15 0.34 -20.69 -21.33 -20.13	1155 460 2699 15 52 74 37 29 80 101 13317 2306 454 10197 111 89 7614 236 283 225 236 1815 5543 251 76 69 8590 6362 116 8791 6190 98 117
						146 147

Table 3 .

 3 2 : Oxygen and multiple sulfur isotope composition of dissolved sulfate ion from the Vena del Gesso gypsum. Error is 1σ.

	Cycle	Facies	Sample name	Δ 33 SSO4 ±0.004‰	δ 34 SSO4 ±0.2‰	δ 18 OSO4 ±0.2‰
	1	Massive selenite	Ma I-01	0.033	21.78	8.24
	2	Massive selenite	MC II-02	0.037	22.65	9.26
	2	Massive selenite	Ma II-07	0.033	22.10	n.d
	3	Massive selenite	Ma III-01	0.029	22.06	8.67
	4	Massive selenite	Ma IV-02	0.029	22.39	8.32
	4	Banded selenite	MT IV-05	0.033	22.15	8.01
	5	Massive selenite	Ma V-03	0.028	21.17	8.43
	5	Banded selenite	Ba V-06	0.037	21.83	8.53
	6	Massive selenite	Ma VI-02	0.036	21.92	9.24
	6	Banded selenite	Ma VI-03	0.034	22.13	8.88
	6	Branching selenite	Br VI-04	0.037	23.05	8.80
	7	Laminar selenite	La VII-02	0.033	21.38	n.d
	7	Massive selenite	Ma VII-03	0.043	23.68	n.d
	7	Banded selenite	Ba VII-04	0.038	21.85	9.05
	7	Branching selenite	Br VII-05	0.023	23.47	8.66
	7	Branching selenite	Br VII-06	0.026	20.73	8.76

Table 4 .

 4 1 Inventory of the samples studied in this work describing the analyses performed on each samples for sites134, 374 and 376.

			Top	Bottom
	Exp. Site Core	Interval	Interval	Lithofacies	mbsf	Sample label	SEM δ 34 SFeS
			(cm)	(cm)
	13 134	7	135	137	dolomitic/pyritic marl laminated with anhydrite	324,35 DSDP-134-7		x
	13 134	8	97	99	anhydrite	344,95 DSDP-134-8	x	x
	13 134 10	81	85	anhydrite	359,7 DSDP-134-10A		x
	13 134 10	110	112	marl	360,1 DSDP-134-10B		x
	42 374 17	35	45	Laminated gypsum UG	411,35 DSDP-374-17A		x
	42 374 17	71	76	marl	411,7 DSDP-374-17B	x	x
	42 374 17	140	144	gypsrudite	412,4 DSDP-374-17C		x
	42 374 18	120	124	porfiroblastic gypsarenite	417,2	DSDP-374-18	x	x
	42 374 19	35	36	dolomitic mudstone	418,35 DSDP-374-19A	x	x
	42 374 19	75	80	porfiroblastic gypsarenite	418,7 DSDP-374-19B		x
	42 374 19	110	111	laminar anhydrite	419,05 DSDP-374-19C	x	x
	42 374 20	107	110	porfiroblastic gypsarenite	421,05 DSDP-374-20A		x
	42 374 20	130	135	Laminated gypsum	421,3 DSDP-374-20B		x
	42 374 21	137	140	massive anhydrite	426,35 DSDP-374-21	x	x
	42 376 16	50	52	marl	140,95 DSDP-376-16A		x
	42 376 16	92	97	gypsum silitite	141,4 DSDP-376-16B		x
	42 376 16	117	121	gypsrudite	141,65 DSDP-376-16C		x
	42 376 16	143	148	laminated gypsum	141,9 DSDP-376-16D		x
	42 376 17	115	117	marl	151,1 DSDP-376-17A		x
	42 376 17	140	145	gypsum silitite	151,4 DSDP-376-17B		x
	42 376 18	115	125	gypsrudite	160,65 DSDP-376-18A	x	x
	42 376 18	141	144	gypsum silitite	160,9 DSDP-376-18B		x
	42 376 19	117	119	gypsrudite	170,15 DSDP-376-19A		x
	42 376 19	144	150	laminated gypsum	170,4 DSDP-376-19B	x	x
	42 376 20	66	70	porfiroblastic gypsarenite	185,6 DSDP-376-20A	x	x
	42 376 20	78	83	laminated gypsum	185,75 DSDP-376-20B	x	x
	42 376 20	141	145	gypsum silitite	186,35 DSDP-376-20C	x	x
	42 376 22	108	110	entherolithic anhydrite	206	DSDP-376-22	x	x

Table 4 .

 4 2 Inventory of the samples studied in this work describing the analyses performed on each samples for site 654.

			Top	Bottom
	Exp. Site Core	Interval	Interval	Lithofacies	mbsf	Sample label SEM δ 34 SFeS
			(cm)	(cm)
	107 654	28	19	24	Laminated gypsum	252,2 IODP-654-28A	x	x
	107 654	28	92	94	marl-grey	252,85 IODP-654-28B
	107 654	28	78	79	marl-dark grey	255,7 IODP-654-28C		x
	107 654	29	140	144	Laminated gypsum	261,835 IODP-654-29A	x
	107 654	29	90	92	marl-brown/yellowish	264,1 IODP-654-29B
	107 654	29	7	8	marl-dark grey	264,75 IODP-654-29C		x
	107 654	30	129	134	Laminated gypsum	272,55	IODP-654-30	x	x
	107 654	31	55	56	marl-brown/yellowish	276,75 IODP-654-31A
	107 654	31	138	143	Laminated gypsum	277,65 IODP-654-31B		x
	107 654	32	124	125	marl	282,05 IODP-654-32A		x
	107 654	32	31	36	Laminated gypsum	282,65 IODP-654-32B		x
	107 654	33	34	35	marl-grey	287,9	IODP-654-33		x
	107 654	34	6	8	marl-very dark grey	290,6 IODP-654-34A	x	x
	107 654	34	109	113	Laminated gypsum	291,65 IODP-654-34B		x
	107 654	34	43	46	Laminated gypsum	294	IODP-654-34C		x
	107 654	35	24	25	marl-dark grey	300,45	IODP-654-35		x
	107 654	36	100	105	Laminated gypsum	310,95 IODP-654-36A		x
	107 654	36	9	10	marl-very dark grey	311,55 IODP-654-36B	x	x

  ). Laminated gypsum samples 34 S=20.8±0.8‰ and ∆ 33 S=0.023±0.006‰) from site 654 have on average lower values than the rest of the dataset. Still, as a whole, samples measured in this study, i.e from deep basins of the Mediterranean, display the same range of δ 34 S isotopic compositions than upper gypsum measured from marginal basins., i.e, sampled in Sicily and Cyprus(Aloisi et al., 2022; Pierre, 1982) and thus slightly lighter values than those measured in PLG gypsum[START_REF] Aloisi | The geochemical riddle of "low-salinity gypsum" deposits[END_REF][START_REF] Evans | Coupled measurements of delta O-18 and delta D of hydration water and salinity of fluid inclusions in gypsum from the Messinian Yesares Member, Sorbas Basin (SE Spain)[END_REF] García-Veigas et al., 2018; Lu et al., 2001) (Figure5. Unraveling the roles of sulfide oxidation and of variable fractionation factors during microbial sulfate reduction ___________________________________________________________________________________________ 224 be due (i) to larger incertitude linked to the extraction and measurement of isotopic composition in the 70-90's compared to nowadays, (ii) the fact that fewer samples were measured in the present study and/or, (iii) that some of the samples considered as primary gypsum precipitates in previous studies were actually secondary gypsum deposits originating from the oxidation of sulfide minerals (these typically occur in marls). We thus consider the measurements performed in our study to reflect a closer approximation of the sulfate isotopic composition of the primary gypsum precipitates deposited in deep Mediterranean basins and we will not consider further the values presented in DSDP/ODP initial reports in this work.

	PLG Chapter 5-: Figure 5. 7 Comparison between δ 34 S measured in this study and presented in ODP/DSDP initial UG marginal UG cores (this study) UG cores (old reports) δ 34 S (VS CDT ‰) 14 16 18 20 22 24 26 Site 134 Site 654 Site 374 Site 376 δ 34 S (VS CDT ‰) 18 19 20 21 22 23 24 UG deep basins DSDP/IODP (this study) DSDP/IODP (old reports)
	report for the same drilling cores.
	Figure 5. 6 δ 34 S of the sulfate ion in equilibrium with PLG (yellow) and UG (green) deposits
	formed in marginal basins of the Mediterranean and of the UG from ODP/DSDP drilling campaings from both this study (blue) and intial drilling reports (orange) 5.6.1.2 Oxygen isotopic composition of sulfate of gypsum
	Different from the sulfur isotopic composition of sulfate, the oxygen isotopic composition

6

). However, the same sections than measured in this work have already been analysed for ODP/DSDP initial reports in the 70'-90's. They generally have significantly lower and more variable δ 34 S isotopic than our dataset

(Figure 5. 6)

. This is particularly striking when looking at the same individual sites that were sampled and measured both in this study and for past IODP/DSDP reports

(Figure 5. 7)

. This might of sulfate measured in upper gypsum from the deep Mediterranean basin ( δ 18 O=12.3±1.7‰, n=27) differs significantly from that of Messinian oceanic sulfate Chapter 5-: Unraveling the roles of sulfide oxidation and of variable fractionation factors during microbial sulfate reduction ___________________________________________________________________________________________ 225 δ18 O=9.8±0.9‰)

  Values measured at site 654 (Tyrrhenian basin) are coherent with the ones found by Waldeck et al., 2022 once corrected for gypsum precipitation (see methods section 2.3.3 for more details). Calcium sulfate deposits from sites located in the eastern Mediterranean basin have intermediate values between MSC ocean/ PLG gypsum and UG from western basins, with values of 11±0.7‰ at site 374 in the ionian abyss and of 12±1.8‰ at site 376 on the Florence Rise western Cyprus.

							Western Med.
		16
		14									Eastern Med.
	O (VS SMOW ‰) δ 18	10 12 8									PLG
											UG
		6									MSC ocean
		4	)	Piedmont basin (6)	Caltanissetta basin (1,4)	Caltanissetta basin (4,6)	Sorbas basin (3)	Tyrrhenian basin (1)	Tyrrhenian basin (This study)	SW Sardinia (This study)	Ionian basin (This study)	W. Cyprus (This study)

___________________________________________________________________________________________ et al., 2022). 5.6.2 Sulfur isotopic composition of iron sulfides extracted from calcium sulfate evaporites and marl interbeds

  from gypsum have a large range of variation in both δ 34 S and ∆ 33 S, with 46 and 0.141‰ of variation respectively. This is particularly striking when compared to the fluctuations of gypsum sulfate isotopic composition from which they were extracted that vary of only 7 and 0.031‰ in δ 34 S and ∆ 33 S respectively. In addition, iron sulfide within evaporite lithologies are present in low concentration (235±352 ppm) and have, on average, particularly high δ 34 S values and low ∆ 33 S values with averages of -7±14 and 0.071±0.028 ‰, respectively (Figure5. 9). Conversely, non-evaporitic marls and dolomitic mudstone interbeds display greater concentrations (3916±4957ppm), less variable isotopic compositions, lower δ 34 S=-28 ±8‰ and higher ∆ 33 S=0.11 ±0.025‰ on average. Except from the difference of repartition of δ 34 S and ∆ 33 S isotopic compositions between iron sulfide extracted from marls and those

					0,2
	S (VS CDT‰ )
	∆ 33
	-50	-40	-30	-20	-10	0	10	20
	Anydrite and rehydrated
	gypsum
	Laminated gypsum		δ 34 S FeSx (VS CDT ‰)
	Clastic gypsum

___________________________________________________________________________________________

Iron sulfides extracted

  Unraveling the roles of sulfide oxidation and of variable fractionation factors during microbial sulfate reduction ___________________________________________________________________________________________ 239 contained higher amount of TOC than measured in their study(Bryant et al., 2020). Therefore, they assume that near equilibrium fractionation (≈-70‰) is the rule during MSR in natural marine environments. Nevertheless, it is difficult to know precisely the environmental conditions under which such ancient deposits formed. Further, factors as the lability of organic matter or nutrients limitation could have maintained low csSRR[START_REF] Bradley | Patterns of sulfur isotope fractionation during microbial sulfate reduction[END_REF], explaining the large offset between sulfide and coeval sulfate observed in the Bryant et al.

	(2020) study.
	Conversely, in the present day, SO4 2-rich Aha Resevoir in South West China, Mengdi et al.

. Whereas they relate pyrite isotopic fluctuations to variations in the rate of microbial consumption in the Peruvian OMZ sediments, Pasquier and collaborators argue that variable fractionation factors during MSR cannot have controlled the variations they observed. They argue that maximal fractionation has been observed in pyrite formed in a 100Myrs old sediment that

Chapter 5-:

(2022) demonstrated that seasonal variations in microbial communities controlled variations of the isotopic fractionation expressed during MSR with -32<εMSR<3‰. This implies that variations in MSR induced fractionation (εMSR) only could control shift in the isotopic composition of preserved iron sulfides, whereas it is often discarded as a control for sedimentary sulfide isotopic composition

[START_REF] Jørgensen | Sulfur biogeochemical cycle of marine sediments[END_REF][START_REF] Masterson | Interpreting multiple sulfur isotope signals in modern anoxic sediments using a full diagenetic model (California-Mexico margin: Alfonso Basin)[END_REF] Pasquier et al., 2021b)

. Nonetheless, it remains unclear if variable fractionation factors generated during MSR could be preserved in the sedimentary record. Further, the range of isotopic fractionation reported in the workd of

Mengdi et al. (2022)

as well as those reported from another hypersaline environment

(Avrahamov et al., 2014)

is smaller than the range of εnet calculated from our samples (between -9 and -60‰). Therefore, it is difficult to know if the entire range of isotopic values measured here could be explained by variations in the fractionation factors during MSR.

Fluctuation in csSRR have been suggested to be a factor that influence both (i) the pooling of sulfur isotopes within the porewater through a faster consumption of sulfate than its replenishment by diffusion

(Pasquier et al., 2021b)

(ii) the modulation of fractionation factors expressed during MSR, with smaller fractionation factor (and therefore relatively enriched Chapter 5-: Unraveling the roles of sulfide oxidation and of variable fractionation factors during microbial sulfate reduction ___________________________________________________________________________________________ 240

  E)Where34 ƐMeasured and 33 λMeasured represent the extent of fractionation expressed during microbial sulfate reduction that directly depends on the cellular specific rate of sulfate reduction and on 34 ƐMin , 33 λMin , 34 ƐMax and 33 λMax that are the minimum and maximum empirical fractionations limits and 34 kmsr and 33 kmsr, pseudo first order rate constants obtained by fitting

equations (D) and (E) to the experimental data of

Sim et al. (2011a)

;

Sim et al. (2011b)

.

isotope composition of UG deposits (stage 3) compared to measured values.

  Data points represent the isotopic composition of iron sulfides preserved in calcium sulfate intervals and non-evaporitic interbeds. Individual curves correspond to the modelled isotopic composition of sulfide resulting from the production of hydrogen sulfide with variable fractionation factors 34 ƐMSR and its consumption by various sulfide reoxidative processes fox. Panel A corresponds to abiotic sulfide oxidation, panel B to phototrophic oxidation and panel C to chemolithotrophic sulfide oxidation. Note

	Chapter 6-Conclusion and perspectives
	___________________________________________________________________________________________
	Figure 6. 4 Modelled that the fox flux here has a different definition than in Figure 6. 3, here fox =1 meaning that no H2S has
	been consumed by the considered re-oxidative process and fox =0 meaning that all of the H2S produced
	by MSR have been consumed by a re-oxidative process
	263

  Andreetto, F., Flecker, R., Aloisi, G., Mancini, A.M., Guibourdenche, L., de Villiers, S. and Krijgsman, W. (2022a) High-amplitude water-level fluctuations at the end of the Mediterranean Messinian Salinity Crisis: Implications for gypsum formation, connectivity and global climate. Earth and Planetary Science Letters 595, 117767. Avrahamov, N., Antler, G., Yechieli, Y., Gavrieli, I., Joye, S.B., Saxton, M., Turchyn, A.V. and Sivan, O. (2014) Anaerobic oxidation of methane by sulfate in hypersaline groundwater of the Dead Sea aquifer. Geobiology 12, 511-528. Balci, N., Shanks III, W.C., Mayer, B. and Mandernack, K.W. (2007) Oxygen and sulfur isotope systematics of sulfate produced by bacterial and abiotic oxidation of pyrite. Geochimica et Cosmochimica Acta 71, 3796-3811. -Barsanti, M., Delbono, I., Schirone, A., Langone, L., Miserocchi, S., Salvi, S. and Delfanti, R. (2011) Sediment reworking rates in deep sediments of the Mediterranean Sea. Sci Total Environ 409, 2959-2970. -Basso, D., Thomson, J. and Corselli, C. (2004) Indications of low macrobenthic activity in the deep sediments of the Eastern Mediterranean Sea. Scientia Marina 68, 53-62. -Beaudoin, G., Taylor, B.E., Rumble, D. and Thiemens, M. (1994) Variations in the sulfur isotope composition of troilite from the Cañon Diablo iron meteorite. Geochimica et Cosmochimica Acta 58, 4253-4255. -Benson, R.H. (1978) 35. THE PALEOECOLOGY OF THE OSTRACODES OF DSDP LEG 42A, Initial Reports of the Deep Sea Drilling Project. US Government Printing Office Washington, pp. 777-787.

	Main text bibliography Main text bibliography
	-Andreetto, F., Mancini, A., Flecker, R., Gennari, R., Lewis, J., Lozar, F., Natalicchio, M.,
	Sangiorgi, F., Stoica, M., Dela Pierre, F. and Krijgsman, W. (2022b) Multi-proxy
	investigation of the post-evaporitic succession of the Piedmont Basin (Pollenzo
	section, NW Italy): A new piece in the Stage 3 puzzle of the Messinian Salinity Crisis.
	Palaeogeography, Palaeoclimatology, Palaeoecology 594, 110961.
	-Andreetto, F., Matsubara, K., Beets, C., Fortuin, A., Flecker, R. and Krijgsman, W.
	(2021b) High Mediterranean water-level during the Lago-Mare phase of the Messinian
	Salinity Crisis: insights from the Sr isotope records of Spanish marginal basins (SE
	Spain).
	-Babel, M. (2007) Depositional environments of a salina-type evaporite basin recorded
	in the Badenian gypsum facies in the northern Carpathian Foredeep.
	-Bąbel, M. and Schreiber, B. (2014) Geochemistry of Evaporites and Evolution of
	Seawater. Treatise on Geochemistry. Ed. by Heinrich D. Holland and Karl K. Turekian.
	Oxford.

, Vasiliev, I. and Krijgsman, W. (2021a)

Freshening of the Mediterranean Salt Giant: controversies and certainties around the terminal (Upper Gypsum and Lago-Mare) phases of the Messinian Salinity Crisis. Earth-Science Reviews 216, 103577. ___________________________________________________________________________________________ 274 --Antler, G., Mills, J.V., Hutchings, A.M., Redeker, K.R. and Turchyn, A.V. (2019) The Sedimentary Carbon-Sulfur-Iron Interplay -A Lesson From East Anglian Salt Marsh Sediments. Frontiers in Earth Science 7. -Antler, G., Turchyn, A.V., Rennie, V., Herut, B. and Sivan, O. (2013) Coupled sulfur and oxygen isotope insight into bacterial sulfate reduction in the natural environment. Geochimica et Cosmochimica Acta 118, 98-117. -Aref, M. (2003) Lithofacies characteristics, depositional environment and karstification of the Late Miocene (Messinian) gypsum deposits in the Northern Western Desert, Egypt. Sedimentology of Egypt 11, 9-27. -Asprey, L.B. (1976) The preparation of very pure fluorine gas. Journal of Fluorine Chemistry 7, 359-361. -Au Yang, D., Landais, G., Assayag, N., Widory, D. and Cartigny, P. (2016) Improved analysis of micro-and nanomole-scale sulfur multi-isotope compositions by gas source isotope ratio mass spectrometry. Rapid Commun Mass Spectrom 30, 897-907. -Avrahami, M. and Golding, R.M. (1968) The oxidation of the sulphide ion at very low concentrations in aqueous solutions. Journal of the Chemical Society A: Inorganic, Physical, Theoretical, 647-651. --Bak, F. and Cypionka, H. (1987) A novel type of energy metabolism involving fermentation of inorganic sulphur compounds. Nature 326, 891-892. ___________________________________________________________________________________________ 275 --Berg, J.S., Duverger, A., Cordier, L., Laberty-Robert, C., Guyot, F. and Miot, J. (2020) Rapid pyritization in the presence of a sulfur/sulfate-reducing bacterial consortium. Scientific Reports 10, 8264. -Berner, R.A. (1982) Burial of organic carbon and pyrite sulfur in the modern ocean: its geochemical and environmental significance. Am. J. Sci.;(United States) 282.

  Blanc, P.-L. (2006) Improved modelling of the Messinian Salinity Crisis and conceptual implications. Palaeogeography, Palaeoclimatology, Palaeoecology 238, 349-372.-Blanc-Valleron, M.-M., Rouchy, J.-M.,Pierre, C., Badaut-Trauth, D. and Schuler, M. (1998) 34. EVIDENCE OF MESSINIAN NONMARINE DEPOSITION AT SITE 968 (CYPRUS LOWER SLOPE), Proceedings of the Ocean Drilling Program, Scientific Results. ODP Sci. Results 160, Texas, USA, pp. 437-445. -Blanc-Valleron, M.M., Pierre, C., Caulet, J.P., Caruso, A., Rouchy, J.M., Cespuglio, G., Sprovieri, R., Pestrea, S. and Di Stefano, E. (2002) Sedimentary, stable isotope and micropaleontological records of paleoceanographic change in the Messinian Tripoli Formation (Sicily, Italy). Palaeogeography, Palaeoclimatology, Palaeoecology 185, 255-286. -Blonder, B., Boyko, V., Turchyn, A.V., Antler, G., Sinichkin, U., Knossow, N., Klein, R. and Kamyshny, A. (2017) Impact of Aeolian Dry Deposition of Reactive Iron Minerals on Sulfur Cycling in Sediments of the Gulf of Aqaba. Frontiers in Microbiology 8. Bourcart, J. (1959) Le plateau continental de la Méditerranée occidentale. C.R Acad. Sci., 249, 1380-1382. -Bourcart, J. (1960) La Méditerranée etla révolution du Pliocène. L'évolution paléogéographique et structurale des domaines méditerranées et alpins d'Europe, in: Fr, M.S.G. (Ed.), Livre à la mémoire du Professeur P.Fallot, pp. 103-118. -Bourillot, R., Vennin, E., Rouchy, J.-M., Blanc-Valleron, M.-M., Caruso, A. and Durlet, C. (2010) The end of the Messinian Salinity Crisis in the western Mediterranean: Insights from the carbonate platforms of south-eastern Spain. Sedimentary Geology 229, 224-253. -Bowles, M.W., Mogollón, J.M., Kasten, S., Zabel, M. and Hinrichs, K.-U. (2014) Global rates of marine sulfate reduction and implications for sub-sea-floor metabolic activities. Science 344, 889. -Boyko, V., Pätzold, J. and Kamyshny, A., Jr (2021) Iron and sulfur speciation and cycling in the sediments of marine systems located in arid environments: the northern Red Sea. Journal of the Geological Society 179. The past 36 million years from fluid inclusions in marine halite. American Journal of Science 313, 713-775. -Briggs, L.I. (1958) Evaporite facies. Journal of Sedimentary Research 28, 46-56. -Brüchert, V., Currie, B., Peard, K., Lass, H., Endler, R., Dübecke, A., Julies, E., Leipe, T. and Zitzmann, S. (2006) "Biogeochemical and physical control on shelf anoxia and water column hydrogen sulphide in the Benguela coastal upwelling system off Namibia," pp. 161-193. -Brüchert, V., Knoblauch, C. and Jørgensen, B.B. (2001) Controls on stable sulfur isotope fractionation during bacterial sulfate reduction in Arctic sediments. Geochimica et Cosmochimica Acta 65, 763-776. Brunner, B. and Bernasconi, S. (2005) A revised isotope fractionation model for dissimilatory sulfate reduction in sulfate reducing bacteria. . Geochimica et Cosmochimica Acta 69, 4759-4771. -Brunner, B., Bernasconi, S.M., Kleikemper, J. and Schroth, M.H. (2005) A model for oxygen and sulfur isotope fractionation in sulfate during bacterial sulfate reduction processes. Geochimica et Cosmochimica Acta 69, 4773-4785.-Bryant, R., Jones, C., Raven, M., Owens, J. and Fike, D. (2020) Shifting modes of iron sulfidization at the onset of OAE-2 drive regional shifts in pyrite δ34S records. Chemical Geology 553, 119808.-Bufe, A., Hovius, N., Emberson, R., Rugenstein, J.K.C., Galy, A., Hassenruck-Gudipati, H.J. and Chang, J.-M. (2021) Co-variation of silicate, carbonate and sulfide weathering drives CO2 release with erosion. Nature Geoscience 14, 211-216.-Bulian, F., Kouwenhoven, T.J., Jiménez-Espejo, F.J., Krijgsman, W., Andersen, N. and Sierro, F.J. (2022) Impact of the Mediterranean-Atlantic connectivity and the late Miocene carbon shift on deep-sea communities in the Western Alboran Basin. Palaeogeography, Palaeoclimatology, Palaeoecology 589, 110841. Monosulfide Oxidation in Natural Sediments: Resolving Microbially Mediated S Transformations Using XANES, Electron Microscopy, and Selective Extractions. Environmental Science & Technology 43, 3128-3134. -Butiseaca, G.-A., Van der Meer, M., Kontakiotis, G., Agiadi, K., Thivaiou, D., Besiou, E., Antonarakou, A., Mulch, A. and Vasiliev, I. (2022) Multiple crises preceded the Mediterranean Salinity Crisis: Aridification and vegetation changes revealed by biomarkers and stable isotopes. Global and Planetary Change 217, 103951. Calmels, D., Gaillardet, J.r.m., Brenot, A.s. and France-Lanord, C. (2007) Sustained sulfide oxidation by physical erosion processes in the Mackenzie River basin: Climatic perspectives. Geology 35, 1003-1006.-Camerlenghi, A., Del Ben, A., Hübscher, C., Forlin, E., Geletti, R., Brancatelli, G., Micallef, A., Saule, M. and Facchin, L. (2020) Seismic markers of the Messinian salinity crisis in the deep IonianBasin. Basin Research 32, 716-738. Canfield, D.E. (2004) The evolution of the Earth surface sulfur reservoir. American Journal of Science 304, 839-861.-Canfield, D.E.,Habicht, K.S. and Thamdrup, B. (2000) The Archean Sulfur Cycle and the Early History of Atmospheric Oxygen. Science 288, 658-661.

	Main text bibliography Main text bibliography Main text bibliography Main text bibliography
	-Canfield, D. (2001) Biogeochemistry of sulfur isotopes. Reviews in mineralogy and
	geochemistry 43, 607-636.
	-Böttcher, M., Thamdrup, B. and Vennemann, T. (2001) Oxygen and sulfur isotope
	fractionation during anaerobic bacterial disproportionation of elemental sulfur.
	Geochimica et Cosmochimica Acta 65, 1601-1609.
	-Böttcher, M.E., Brumsack, H.-J. and De Lange, G.J. (1998) Sulfate reduction and related
	stable isotope (^ 3^ 4S,^ 1^ 8O) variations in interstitial waters from the Eastern
	Mediterranean, PROCEEDINGS-OCEAN DRILLING PROGRAM SCIENTIFIC RESULTS.
	NATIONAL SCIENCE FOUNDATION, pp. 365-376.
	-Boulton, S.J., Smart, C.W., Consolaro, C. and Snider, A. (2016) The Miocene-Pliocene
	boundary and the Messinian Salinity Crisis in the easternmost Mediterranean: Insights
	from the Hatay Graben (Southern Turkey). Sedimentary Geology 332, 51-67.

___________________________________________________________________________________________ 276 --Boetius, A., Ravenschlag, K., Schubert, C.J., Rickert, D., Widdel, F., Gieseke, A., Amann, R.,

Jørgensen, B.B., Witte, U. and Pfannkuche, O. (2000)

A marine microbial consortium apparently mediating anaerobic oxidation of methane. Nature 407, 623-626.

-Borrelli, M., Perri, E., Critelli, S. and Gindre-Chanu, L. (

2021

) The onset of the Messinian Salinity Crisis in the central Mediterranean recorded by pre-salt carbonate/evaporite deposition. Sedimentology 68, 1159-1197. -Borsetti, A.M., Curzi, P.V., Landuzzi, V., Ricci Lucchi, F., Sartori, R., Tomadin, L. and Zuffa, G.G. (1990) Messinian and pre-messinian sediments from ODP Leg 107, Sites 652 and 654 in the Tyrrhenian Sea: sedimentologic and petrologic study and possible comparison with Italian sequences. Proceedings of the ODP, Scientific Results 107, 169-186. -Boschetti, T., Cortecchi, G., Toscani, L. and Iacumin, P. (2011) Sulfur and oxygen isotope compositions of Upper Triassic sulfates from Northerm Apennines (Italy): palaeogeographic and hidrogeochemical implications. Geologica Acta, 129-147. -Bourcart, J. (1958) Morphologie du Précontient des Pyrénées à Sardaigne avec une carte au 1/200.000, LXXXIII Colloque Intern. du CNRS. Topographie et Géologie des Profondeurs Océaniques. ___________________________________________________________________________________________ 277 --Bradley, A.S., Leavitt, W.D., Schmidt, M., Knoll, A.H., Girguis, P.R. and Johnston, D.T. (2016) Patterns of sulfur isotope fractionation during microbial sulfate reduction. Geobiology 14, 91-101. -Bradley, J., Arndt, S., Amend, J., Burwicz, E., Dale, A.W., Egger, M. and LaRowe, D.E. (2020) Widespread energy limitation to life in global subseafloor sediments. Science advances 6, eaba0697. -Bradley, J.A., Arndt, S., Amend, J.P., Burwicz-Galerne, E. and LaRowe, D.E. (2022) Sources and Fluxes of Organic Carbon and Energy to Microorganisms in Global Marine Sediments. Frontiers in Microbiology 13. -Brand, W.A., Coplen, T., B. , Vogl, J., Rosner, M. and Prohaska, T. (2014) Assessment of international reference materials for isotope-ratio analysis (IUPAC Technical Report). Pure and Applied Chemistry 86, 425-467. -Brennan, S., Lowenstein, T. and Cendón, D. (2013) The major-ion composition of Cenozoic seawater: ___________________________________________________________________________________________ 278 --Bulian, F., Sierro, F.J., Ledesma, S., Jiménez-Espejo, F.J. and Bassetti, M.-A. (2021) Messinian West Alboran Sea record in the proximity of Gibraltar: Early signs of Atlantic-Mediterranean gateway restriction. Marine Geology 434, 106430. -Burke, A., Present, T.M., Paris, G., Rae, E.C.M., Sandilands, B.H., Gaillardet, J., Peucker-Ehrenbrink, B., Fischer, W.W., McClelland, J.W., Spencer, R.G.M., Voss, B.M. and Adkins, J.F. (2018) Sulfur isotopes in rivers: Insights into global weathering budgets, pyrite oxidation, and the modern sulfur cycle. Earth and Planetary Science Letters 496, 168-177. -Burton, E.D., Bush, R.T., Johnston, S.G., Sullivan, L.A. and Keene, A.F. (2011) Sulfur biogeochemical cycling and novel Fe-S mineralization pathways in a tidally re-flooded wetland. Geochimica et Cosmochimica Acta 75, 3434-3451. -Burton, E.D., Bush, R.T., Sullivan, L.A., Hocking, R.K., Mitchell, D.R.G., Johnston, S.G., Fitzpatrick, R.W., Raven, M., McClure, S. and Jang, L.Y. (2009) Iron--Callbeck, C.M., Canfield, D.E., Kuypers, M.M.M., Yilmaz, P., Lavik, G., Thamdrup, B., Schubert, C.J. and Bristow, L.A. (2021) Sulfur cycling in oceanic oxygen minimum zones. Limnology and Oceanography n/a. ___________________________________________________________________________________________ 279 ---Canfield, D.E., Kristensen, E. and Thamdrup, B. (2005) Aquatic geomicrobiology. -Canfield, D.E., Stewart, F.J., Thamdrup, B., De Brabandere, L., Dalsgaard, T., Delong, E.F., Revsbech, N.P. and Ulloa, O. (2010) A Cryptic Sulfur Cycle in Oxygen-Minimum-Zone Waters off the Chilean Coast. Science 330, 1375-1378. -Capella, W., Flecker, R., Hernández-Molina, F.J., Simon, D., Meijer, P.T., Rogerson, M., Sierro, F.J. and Krijgsman, W. (2019) Mediterranean isolation preconditioning the Earth System for late Miocene climate cooling. Scientific Reports 9, 3795.

  ) Oxidative transformation of iron monosulfides and pyrite in estuarine sediments: Implications for trace metals mobilisation. Journal of Environmental Management 186, 158-166.-CIESM, C.(2008) The Messinian Salinity Crisis from mega-deposits to microbiology-a consensus report, CIESM Workshop monographs. Ciesm Monaco, pp. 1-168. Proceedings of the Ocean Drilling Program, Initial Reports, v. 161: College Station, Texas, Ocean Drilling Program. -Cornée, J.-J., Saint Martin, J.-P., Conesa, G., Münch, P., André, J.-P., Saint Martin, S. and Roger, S. (2004) Correlations and sequence stratigraphic model for Messinian carbonate platforms of the western and central Mediterranean. International Journal of Earth Sciences 93, 621-633. Damsté, S.J.S., Frewin, N.L., Kenig, F. and De Leeuw, J.W. (1995) Molecular indicators for palaeoenvironmental change in a Messinian evaporitic sequence (Vena del Gesso, Italy). I: Variations in extractable organic matter of ten cyclically deposited marl beds. Organic Geochemistry 23, 471-483. -Dauphas, N. and Schauble, E.A. (2016) Mass Fractionation Laws, Mass-Independent Effects, and Isotopic Anomalies. Annual Review of Earth and Planetary Sciences 44, 709-783. -de Deckker, P., Chivas, A.R. and Shelley, J.M.G. (1988) Paleoenvironment of the Messinian Mediterranean "Lago Mare" from Strontium and Magnesium in Ostracode Shells. PALAIOS 3, 352-358. -de Lange, G.J. and Krijgsman, W. (2010) Messinian salinity crisis: A novel unifying shallow gypsum/deep dolomite formation mechanism. Marine Geology 275, 273-277. Flocculent layers and bacterial mats in the mudstone interbeds of the Primary Lower Gypsum unit (Tertiary Piedmont basin, NW Italy): Archives of palaeoenvironmental changes during the Messinian salinity crisis. Marine Geology 355, 71-87. Devereux, R., Lehrter, J.C., Cicchetti, G., Beddick, D.L., Yates, D.F., Jarvis, B.M., Aukamp, J. and Hoglund, M.D. (2019) Spatially variable bioturbation and physical mixing drive the sedimentary biogeochemical seascape in the Louisiana continental shelf hypoxic zone. Biogeochemistry 143, 151-169. -Dickens, G., Fewless, T., Thomas, E., Bralower, T., Gingerich, L. and Schmitz, P. (2003) Excess barite accumulation during the Paleocene-Eocene Thermal Maximum: Massive input of dissolved barium from seafloor gas hydrate reservoirs. Spec. Pap. Geol. Soc. Am. 369. -Dinelli, E., Testa, G., G, C. and Barbieri, M. (1999) Stratigraphic and petrographic constraints to trace element and isotope geochemistry of Messinian sulfates of Tuscany. Memorie della Societa geologica Italiana 54, 61-74. -Drinia, H., Antonarakou, A., Tsaparas, N. and Kontakiotis, G. (2007) Palaeoenvironmental conditions preceding the Messinian Salinity Crisis: A case study from Gavdos Island. Geobios 40, 251-265.

	Main text bibliography Main text bibliography
	___________________________________________________________________________________________ ___________________________________________________________________________________________
	Claypool continued: Extending the isotopic record of sedimentary sulfate. Chemical anaerobic oxidation of methane is dependent on methane concentration. Earth and
	Geology 513, 200-225. Planetary Science Letters 399, 61-73.
	-Dela Pierre, F., Natalicchio, M., Ferrando, S., Giustetto, R., Birgel, D., Carnevale, G.,
	Gier, S., Lozar, F., Marabello, D. and Peckmann, J. (2015) Are the large filamentous
	-Costanzo, A., Cipriani, M., Feely, M., Cianflone, G. and Dominici, R. (2019) Messinian microfossils preserved in Messinian gypsum colorless sulfide-oxidizing bacteria?
	twinned selenite from the Catanzaro Trough, Calabria, Southern Italy: field, Geology 43, 855-858.
	petrographic and fluid inclusion perspectives. Carbonates and Evaporites 34, 743-756. -Deroo, G. (1978) Organic geochemistry of some Neogene cores from Sites 374, 375,
	377, and 378: Leg 42A, eastern Mediterranean Sea.
	281 282

-Cita, M. and

Ryan, W. (1978a)

Messinian paleoenvironments.

-Cita, M. and

Ryan, W.B.F. (1978b)

Messinian paleoenvironments.

-Cita, M.B., Santambrogio, S.,

[START_REF] Cita | Messinian Paleoenvironments: New Evidence from the Tyrrhenian Sea (ODP LEG 107)[END_REF]

14. MESSINIAN PALEOENVIRONMENTS: NE W EVIDENCE FRO M TH E TYRRHENIA N SEA (ODP LEG 107). -Clauzon, G. (1974) L'hypothèse eustatique et le creusement prépliocène de la vallée du Rhône, Annales de Géographie. JSTOR, pp. 129-140. -Clauzon, G., Suc, J.-P., Gautier, F., Berger, A. and Loutre, M.-F. (1996) Alternate interpretation of the Messinian salinity crisis: Controversy resolved? Geology 24, 363. -Claypool, G.E., Holser, W.T., Kaplan, I.R., Sakai, H. and Zak, I. (1980) The age curves of sulfur and oxygen isotopes in marine sulfate and their mutual interpretation. Chemical Geology 28, 199-260. -Cohen, A. (1993) Halite-clay interplay in the Israeli Messinian. Sedimentary Geology 86, 211-228. -Colangelo-Lillis, J., Pelikan, C., Herbold, C.W., Altshuler, I., Loy, A., Whyte, L.G. and Wing, B.A. (2019) Diversity decoupled from sulfur isotope fractionation in a sulfatereducing microbial community. Geobiology 17, 660-675. -Comas, M., Zahn, R. and Klaus, A. (1996) -Crockford, P.W., Kunzmann, M., Bekker, A., Hayles, J., Bao, H., Halverson, G.P., Peng, Y., Bui, T.H., Cox, G.M., Gibson, T.M., Wörndle, S., Rainbird, R., Lepland, A., Swanson-Hysell, N.L., Master, S., Sreenivas, B., Kuznetsov, A., Krupenik, V. and Wing, B.A. (2019) --Decima, A. (1964) Ostracodi del gen. Cyprideis Jones del Neogene e del Quaternario italiani. -Decima, A. and Wezel, F. (1973) LATE MIOCENE EVAPORITES OF THE CENTRAL SICILIAN BASIN, ITALY.

-Dela Pierre, F., Bernardi, E., Cavagna, S., Clari, P., Gennari, R., Irace, A., Lozar, F., Lugli, S., Manzi, V., Natalicchio, M.,

Roveri, M. and Violanti, D. (2011)

The record of the Messinian salinity crisis in the Tertiary Piedmont Basin (NW Italy): The Alba section revisited. Palaeogeography, Palaeoclimatology, Palaeoecology 310, 238-255. -Dela Pierre, F., Clari, P., Bernardi, E., Natalicchio, M., Costa, E., Cavagna, S., Lozar, F., Lugli, S., Manzi, V., Roveri, M. and Violanti, D. (2012) Messinian carbonate-rich beds of the Tertiary Piedmont Basin (NW Italy): Microbially-mediated products straddling the onset of the salinity crisis. Palaeogeography, Palaeoclimatology, Palaeoecology 344-345, 78-93. -Dela Pierre, F., Clari, P., Natalicchio, M., Ferrando, S., Giustetto, R., Lozar, F., Lugli, S., Manzi, V., Roveri, M. and Violanti, D. (2014) -Deusner, C., Holler, T., Arnold, G.L., Bernasconi, S.M., Formolo, M.J. and Brunner, B. (2014) Sulfur and oxygen isotope fractionation during sulfate reduction coupled to --Duggen, S., Hoernle, K., van den Bogaard, P. and Harris, C. (2004) Magmatic evolution of the Alboran region: The role of subduction in forming the western Mediterranean and causing the Messinian Salinity Crisis. Earth and Planetary Science Letters 218, 91-108.

  Flecker, R., Krijgsman, W., Capella, W., de Castro Martins, C., Dmitrieva, E., Mayser, J.P., Marzocchi, A., Modestu, S., Lozano, D.O. and Simon, D. (2015) Evolution of the Late Miocene Mediterranean-Atlantic gateways and their impact on regional and global environmental change.-Fontes, J.C., Letolle, R.,Nesteroff, W. and Ryan, W. (1973) Oxygen, Carbon, Sulfur, and Hydrogen Stable Isotopes in Carbonate and Sulfate Mineral Phases of Neogene Evaporites, Sediments, and in Interstitial Waters. Deep Sea Drill. Proj. 13. -Forrest, J. and Newman, L. (1977) Silver-110 microgram sulfate analysis for the short time resolution of ambient levels of sulfur aerosol. Analytical Chemistry 49, 1579-1584. -Fossing, H. and Jørgensen, B.B. (1989) Measurement of bacterial sulfate reduction in sediments: Evaluation of a single-step chromium reduction method. Biogeochemistry 8, 205-222. -Fossing, H. and Jørgensen, B.B. (1990) Oxidation and reduction of radiolabeled inorganic sulfur compounds in an estuarine sediment, Kysing Fjord, Denmark. Geochimica et Cosmochimica Acta 54, 2731-2742. -Fotherby, A., Bradbury, H.J., Antler, G., Sun, X., Druhan, J.L. and Turchyn, A.V. (2021) Modelling the Effects of Non-Steady State Transport Dynamics on the Sulfur and Oxygen Isotope Composition of Sulfate in Sedimentary Pore Fluids. Frontiers in Earth Science 8. Garcia-Castellanos, D., Estrada, F., Jiménez-Munt, I., Gorini, C., Fernàndez, M., Vergés, J. and De Vicente, R. (2009) Catastrophic flood of the Mediterranean after the Messinian salinity crisis. Nature 462, 778-781. Garrison, R., BC, S. and FH, F. (1978) Sedimentary petrology and structures of Messinian evaporitic sediments in the Mediterranean Sea, leg 42A, Deep Sea Drilling Project. -Gass, I. and Cockbain, A. (1961) Notes on the occurrence of gypsum in Cyprus. HM Stationery Office. -Geletti, R., Zgur, F., Del Ben, A., Buriola, F., Fais, S., Fedi, M., Forte, E., Mocnik, A., Paoletti, V., Pipan, M., Ramella, R., Romeo, R. and Romi, A. (2014) The Messinian Salinity Crisis: New seismic evidence in the West-Sardinian Margin and Eastern Sardo-Provençal basin (West Mediterranean Sea). Marine Geology 351, 76-90. -Geng, L., Savarino, J., Savarino, C.A., Caillon, N., Cartigny, P., Hattori, S., Ishino, S. and Yoshida, N. (2018) A simple and reliable method reducing sulfate to sulfide for multiple sulfur isotope analysis. Rapid Commun Mass Spectrom 32, 333-341. -Gennari, R., Lozar, F., Turco, E., Pierre, F.D., Lugli, S., Manzi, V., Natalicchio, M., Roveri, M., Schreiber, B. and Taviani, M. (2018) Integrated stratigraphy and paleoceanographic evolution of the pre-evaporitic phase of the Messinian salinity crisis in the Eastern Mediterranean as recorded in the Tokhni section (Cyprus island). Newsletters on Stratigraphy 51, 33-55.

	Main text bibliography Main text bibliography
	Kamyshny, A.,
	Jr. (2019) Sulfide oxidation affects the preservation of sulfur isotope signals. Geology -Garcia-Orellana, J., Pates, J.M., Masqué, P., Bruach, J.M. and Sanchez-Cabeza, J.A.
	47, 739-743. (2009) Distribution of artificial radionuclides in deep sediments of the Mediterranean
	-Findlay, A.J. and Kamyshny, A. (2017) Turnover Rates of Intermediate Sulfur Species Sea. Sci Total Environ 407, 887-898.
	(Sx2-, S0, S2O32-, S4O62-, SO32-) in Anoxic Freshwater and Sediments. Frontiers in -García-Veigas, J., Cendón, D.I., Gibert, L., Lowenstein, T.K. and Artiaga, D. (2018)
	Microbiology 8. Geochemical indicators in Western Mediterranean Messinian evaporites: Implications
	for the salinity crisis. Marine Geology 403, 197-214.

-Findlay, A.J.,

Pellerin, A., Laufer, K. and Jørgensen, B.B. (2020)

Quantification of sulphide oxidation rates in marine sediment. Geochimica et Cosmochimica Acta 280, 441-452. ___________________________________________________________________________________________ 284 --Friedman, G. (2007) THIN-SECTION PETROGRAPHY OF THE MEDITERRANEAN EVAPORITES 22 . 1 . PETROGRAPHIC DATA AND COMMENTS ON THE DEPOSITIONAL ENVIRONMENT OF THE MIOCENE SULFATES AND DOLOMITES AT SITES 124 , 132 AND 134 , WESTERN MEDITERRANEAN SEA. -Froelich, P.N., Klinkhammer, G.P., Bender, M.L., Luedtke, N.A., Heath, G.R., Cullen, D., Dauphin, P., Hammond, D., Hartman, B. and Maynard, V. (1979) Early oxidation of organic matter in pelagic sediments of the eastern equatorial Atlantic: suboxic diagenesis. Geochimica et Cosmochimica Acta 43, 1075-1090. -Fry, B., Cox, J., Gest, H. and Hayes, J.M. (1986) Discrimination between 34S and 32S during bacterial metabolism of inorganic sulfur compounds. J Bacteriol 165, 328-330. --Garrels, R.M. and Lerman, A. (1984) Coupling of the sedimentary sulfur and carbon cycles; an improved model. American Journal of Science 284, 989-1007. ___________________________________________________________________________________________ 285 -

  signatures of the surface microbial mat environment in sediments underlying low-oxygen cyanobacterial mats. Geobiology 20, 60-78.-Gröger, J., Franke, J.,Hamer, K. and Schulz, H.D. (2009) Quantitative Recovery of Elemental Sulfur and Improved Selectivity in a Chromium-Reducible Sulfur Distillation. Geostandards and Geoanalytical Research 33, 17-27.-Grossi, F.,Gliozzi, E., Anadón, P., Castorina, F. and Voltaggio, M. (2015) Is Cyprideis agrigentina Decima a good paleosalinometer for the Messinian Salinity Crisis? Morphometrical and geochemical analyses from the Eraclea Minoa section (Sicily). Palaeogeography, Palaeoclimatology, Palaeoecology 419, 75-89.-Guibourdenche, L., Cartigny, P., Dela Pierre, F., Natalicchio, M. and Aloisi, G. (2022) Cryptic sulfur cycling during the formation of giant gypsum deposits. Earth and Planetary Science Letters 593, 117676. -Gvirtzman, Z., Manzi, V., Calvo, R., Gavrieli, I., Gennari, R., Lugli, S., Reghizzi, M. and Roveri, M. (2017) Intra-Messinian truncation surface in the Levant Basin explained by subaqueous dissolution. Geology 45, 915-918. -Habicht, K.S. and Canfield, D.E. (2001) Isotope fractionation by sulfate-reducing natural populations and the isotopic composition of sulfide in marine sediments. Geology 29, 555-558. Harvie, C.E. and Weare, J.H. (1980) The prediction of mineral solubilities in natural waters: the NaKMgCaClSO4H2O system from zero to high concentration at 25° C. Geochimica et Cosmochimica Acta 44, 981-997. -Hay, W.W., Migdisov, A., Balukhovsky, A.N., Wold, C.N., Flögel, S. and Söding, E. (2006) Evaporites and the salinity of the ocean during the Phanerozoic: Implications for climate, ocean circulation and life. Palaeogeography, Palaeoclimatology, Palaeoecology 240, 3-46. -Hayes, J.M. and Waldbauer, J.R. (2006) The carbon cycle and associated redox processes through time. Philos Trans R Soc Lond B Biol Sci 361, 931-950. -He, T., Corso, J.D., Newton, R.J., Wignall, P.B., Mills, B.J.W., Todaro, S., Stefano, P.D., Turner, E.C., Jamieson, R.A., Randazzo, V., Rigo, M., Jones, R.E. and Dunhill, A.M. (2020) An enormous sulfur isotope excursion indicates marine anoxia during the end-Triassic mass extinction. Science Advances 6, eabb6704. -Heida, H., Raad, F., Garcia-Castellanos, D., Jiménez-Munt, I., Maillard, A. and Lofi, J. (2021) Flexural-isostatic reconstruction of the Western Mediterranean during the Messinian Salinity Crisis: Implications for water level and basin connectivity. Basin Research 34, 50-80. Insights from the Valle di Manche section (Lower-Middle Pleistocene, southern Italy). Palaeogeography, Palaeoclimatology, Palaeoecology 585, 110730. -Hsu, K. and Montadert, L. (1978) Initial Reports of the Deep Sea Drilling Project, 42 Pt. 1. Hsü, K., Montadert, L., Bernoulli, D., Bizon, G., Cita, M., Erickson, A., Fabricius, F., Garrison, R., Kidd, R. and Mélières, F. (1978c) Sites 375 and 376: Florence Rise. Initial Reports of the Deep-Sea Drilling Project. Edited by KJ Hsü and L. Montadert 42, 219-304.-Hsü, K.,Montadert, L., Bernoulli, D., Bizon, G., Cita, M., Erickson, A. and Wright, R. (1978d) Site 374: Messina abyssal plain. Initial Rep. Deep Sea Drill. Proj 42, 175-217.-Hsü, K.J.,Ryan, W.B.F. and Cita, M.B. (1973) Late Miocene Desiccation of the Mediterranean. Nature 242, 240-244.-Hsu¨, K.J. (1972) Origin of saline giants: A critical review after the discovery of the Mediterranean Evaporite. Earth-Science Reviews 8, 371-396.-Hübscher, C., Cartwright, J., Cypionka, H., DeLange, G.J., Robertson, A., Suc, J.P. and Urai, J.L. (2007) Global look at salt giants. Eos, Transactions American Geophysical Union 88, 177-179.-Hurtgen, M.,Lyons, T., Ingall, E. and Cruse, A. (1999) Anomalous enrichments of iron monosulfide in euxinic marine sediments and the role of H2S in iron sulfide transformations: Examples from Effingham Inlet, Orca Basin, and the Black Sea. American Journal of Science -AMER J SCI 299, 556-588. Sulfur isotope insights into microbial sulfate reduction: When microbes meet models. Geochimica et Cosmochimica Acta 71, 3929-3947. -Johnston, D.T., Farquhar, J., Habicht, K.S. and Canfield, D.E. (2008) Sulphur isotopes and the search for life: strategies for identifying sulphur metabolisms in the rock record and beyond. Geobiology 6, 425-435. Johnston, D.T., Gill, B.C., Masterson, A., Beirne, E., Casciotti, K.L., Knapp, A.N. and Berelson, W. (2014) Placing an upper limit on cryptic marine sulphur cycling. Nature 513, 530-533. -Jones, D.S. and Fike, D.A. (2013) Dynamic sulfur and carbon cycling through the end-Ordovician extinction revealed by paired sulfate-pyrite δ34S. Earth and Planetary Science Letters 363, 144-155.

	Main text bibliography Main text bibliography Main text bibliography Main text bibliography
	___________________________________________________________________________________________
	-Jørgensen, B., Bang, M. and Blackburn, T.H. (1990) Anaerobic mineralization in marine sediments from the Baltic Sea-North Sea transition. Marine Ecology-progress Series -MAR ECOL-PROGR SER 59, 39-54. -Jørgensen, B.B. (1977) Distribution of colorless sulfur bacteria (Beggiatoa spp.) in a coastal marine sediment. Marine Biology 41, 19-28. -Jørgensen, B.B. (2021) SULFUR BIOGEOCHEMICAL CYCLE OF MARINE SEDIMENTS. isotope compositions preserve -Hardie, L.A. and Eugster, H.P. (1971) THE DEPOSITIONAL ENVIRONMENT OF MARINE Geochemical Perspectives 10, 145-146.
	EVAPORITES: A CASE FOR SHALLOW, CLASTIC ACCUMULATION. Sedimentology 16,
	187-220.
	-Hardie, L.A. and Lowenstein, T. (2004) Did the Mediterranean Sea Dry Out During the -Hsü, K., Montadert, L., Bernoulli, D., Bizon, G., Cita, M., Erickson, A., Fabricius, F.,
	Miocene? A Reassessment of the Evaporite Evidence from DSDP Legs 13 and 42A Garrison, R., Kidd, R. and Melieres, F. (1978a) Site 371: South Balearic Basin. Initial -Johnston, D.T., Farquhar, J., Wing, B.A., Kaufman, A.J., Canfield, D.E. and Habicht, K.S.
	Cores. Journal of Sedimentary Research 74. Reports of the Deep Sea Drilling Project 42, 29-57. (2005) Multiple sulfur isotope fractionations in biological systems: A case study with
	sulfate reducers and sulfur disproportionators. American Journal of Science 305, 645-
	660.
	286

.,

Kinsman-Costello, L., Sheldon, N.D. and Fike, D.A. (2022)

Sedimentary pyrite sulfur -Habicht, K.S.,

Canfield, D.E. and Rethmeier, J.o. (1998)

Sulfur isotope fractionation during bacterial reduction and disproportionation of thiosulfate and sulfite. Geochimica et Cosmochimica Acta 62, 2585-2595. -Halevy, I., Peters, S.E. and Fischer, W.W. (2012) Sulfate Burial Constraints on the Phanerozoic Sulfur Cycle. Science 337, 331-334. -Haq, B., Gorini, C., Baur, J., Moneron, J. and Rubino, J.-L. (2020) Deep Mediterranean's Messinian evaporite giant: How much salt? Global and Planetary Change 184, 103052. -Hardenberg, M.F. and Robertson, A.H.F. (2007) Sedimentology of the NW margin of the Arabian plate and the SW-NE trending Nahr El-Kabir half-graben in northern Syria during the latest Cretaceous and Cenozoic. Sedimentary Geology 201, 231-266. -Hardie, L.A. (1991) On the significance of evaporites. Annual Review of Earth and Planetary Sciences 19, 131. -Harrab, S., Mannai-Tayech, B., Rabhi, M. and Zargouni, F. (2013) Study of a Neogene basin dynamics: The "Bizerte basin", northeastern Tunisia: relevance to the global Messinian Salinity Crisis. Comptes Rendus Geoscience 345, 251-261.

___________________________________________________________________________________________ 287 --Hersey, J. (1965) Sedimentary basins of the Mediterranean Sea, Submarine geology and geophysics. Butterworths London, pp. 75-91. -Hoareau, G., Monnin, C. and Odonne, F. (2011) The stability of gypsum in marine sediments using the entire ODP/IODP porewater composition database. Marine Geology 279, 87-97. -Holser, W.T. and Kaplan, I.R. (1966) Isotope geochemistry of sedimentary sulfates. Chemical Geology 1, 93-135.

-Horita, J., Zimmermann, H. and Holland, H.D. (

2002

) Chemical evolution of seawater during the Phanerozoic: Implications from the record of marine evaporites. Geochimica et Cosmochimica Acta 66, 3733-3756.

-Houghton, J., Scarponi, D., Capraro, L. and Fike, D.A. (2022) Impact of sedimentation, climate and sea level on marine sedimentary pyrite sulfur isotopes:

-Hsü, K.,

Montadert, L., Bernoulli, D., Bizon, G., Cita, M., Erickson, A., Fabricius, F., Garrison, R., Kidd, R. and Mélières, F. (1978b)

Site 378: Cretan Basin. Initial Rep. Deep Sea Drill. Proj 42, 321-357. ___________________________________________________________________________________________ 288 --Iaccarino, S. and Bossio, A. (1999) Paleoenvironment of uppermost Messinian sequences in the western Mediterranean (Sites 974, 975, and 978), Proceedings of the Ocean Drilling Program, Scientific Results. Ocean Drilling Program College Station, TX, pp. 529-541. -Isaji, Y., Kawahata, H., Takano, Y., Ogawa, N.O., Kuroda, J., Yoshimura, T., Lugli, S., Manzi, V., Roveri, M. and Ohkouchi, N. (2019) Diazotrophy Drives Primary Production in the Organic-Rich Shales Deposited Under a Stratified Environment During the Messinian Salinity Crisis (Vena del Gesso, Italy). Frontiers in Earth Science 7. -Jiang, L., Zhao, M., Shen, A., Huang, L., Chen, D. and Cai, C. (2022) Pulses of atmosphere oxygenation during the Cambrian radiation of animals. Earth and Planetary Science Letters 590, 117565. -Johnston, D.T., Farquhar, J. and Canfield, D.E. (2007) ___________________________________________________________________________________________ 289 --Jørgensen, B.B., Findlay, A.J. and Pellerin, A. (2019) The Biogeochemical Sulfur Cycle of Marine Sediments. Frontiers in Microbiology 10. -Jørgensen, B.B. and Kasten, S. (2006) Sulfur Cycling and Methane Oxidation, in: Schulz, H.D., Zabel, M. (Eds.), Marine Geochemistry. Springer Berlin Heidelberg, Berlin, Heidelberg, pp. 271-309.

  Turchyn, A.V., Kasten, S. and Jørgensen, B.B. (2021b) Isotopically "heavy" pyrite in marine sediments due to high sedimentation rates and non-steady-state deposition.Geology 49, 816-821. Luther, G.W., Giblin, A., Howarth, R.W. and Ryans, R.A. (1982) Pyrite and oxidized iron mineral phases formed from pyrite oxidation in salt marsh and estuarine sediments. Geochimica et Cosmochimica Acta 46, 2665-2669.-Lymer, G., Lofi, J., Gaullier, V.,[START_REF] Lymer | The Western Tyrrhenian Sea revisited: New evidence for a rifted basin during the Messinian Salinity Crisis[END_REF] The Western Tyrrhenian Sea revisited: New evidence for a rifted basin during the Messinian Salinity Crisis. Marine Geology 398, 1-21.-Lyons, T.W. (1997) Sulfur isotopic trends and pathways of iron sulfide formation in upper Holocene sediments of the anoxic Black Sea. Geochimica et Cosmochimica Acta 61, 3367-3382. -Lyons, T.W., Reinhard, C.T. and Planavsky, N.J. (2014) The rise of oxygen in Earth's early ocean and atmosphere. Nature 506, 307-315. -Madof, A.S., Bertoni, C. and Lofi, J. (2019) Discovery of vast fluvial deposits provides evidence for drawdown during the late Miocene Messinian salinity crisis. Geology 47, 171-174. 2013) Age refinement of the Messinian salinity crisis onset in the Mediterranean. Terra Nova 25, 315-322. -Manzi, V., Lugli, S., Lucchi, F.R. and Roveri, M. (2005) Deep-water clastic evaporites deposition in the Messinian Adriatic foredeep (northern Apennines, Italy): did the Mediterranean ever dry out? Sedimentology 52, 875-902. Manzi, V., Roveri, M., Gennari, R., Bertini, A., Biffi, U., Giunta, S., Iaccarino, S.M., Lanci, L., Lugli, S., Negri, A., Riva, A., Rossi, M.E. and Taviani, M. (2007) The deep-water counterpart of the Messinian Lower Evaporites in the Apennine foredeep: The Fanantello section (Northern Apennines, Italy). Palaeogeography, Palaeoclimatology, Palaeoecology 251, 470-499. -Marabini, S. and Poluzzi, A. (1977) Le Crisia (Bryozoa, cyclostomata) del Messiniano inferiore della Romagna Occidentale. Giorn. Geol 42, 165-180. -Marin-Carbonne, J., Decraene, M., Havas, R., Remusat, L., Pasquier, V., Alléon, J., Zeyen, N., Bouton, A., Bernard, S. and Escrig, S. (2022) Early precipitated micropyrite in microbialites: A time capsule of microbial sulfur cycling. Geochemical Perspectives Letters 21, 7-12. -Markovic, S., Paytan, A., Li, H. and Wortmann, U.G. (2016) A revised seawater sulfate oxygen isotope record for the last 4Myr. Geochimica et Cosmochimica Acta 175, 239-251.

	Main text bibliography Main text bibliography
	___________________________________________________________________________________________ ___________________________________________________________________________________________
	the Triassic-Jurassic boundary mass extinction. Geochimica et Cosmochimica Acta 231, a further step towards a new stratigraphic framework for Eastern Mediterranean.
	73-87. Basin Research 28, 207-236.
	-Liu, J., Pellerin, A., Antler, G., Kasten, S., Findlay, A.J., Dohrmann, I., Røy, H., Turchyn,
	A.V. and Jørgensen, B.B. (2020a) Early diagenesis of iron and sulfur in Bornholm Basin -Manzi, V., Lugli, S., Roveri, M. and Charlotte Schreiber, B. (2009) A new facies model
	sediments: The role of near-surface pyrite formation. Geochimica et Cosmochimica for the Upper Gypsum of Sicily (Italy): chronological and palaeoenvironmental
	Acta 284, 43-60. constraints for the Messinian salinity crisis in the Mediterranean. Sedimentology 56,
	1937-1960.
	293 294

-Liu, J.,

Pellerin, A., Izon, G., Wang, J., Antler, G., Liang, J., Su, P., Jørgensen, B.B. and Ono, S. (2020b)

The multiple sulphur isotope fingerprint of a sub-seafloor oxidative sulphur cycle driven by iron. Earth and Planetary Science Letters 536, 116165. --Magnuson, E., Mykytczuk, N.C.S., Pellerin, A., Goordial, J., Twine, S.M., Wing, B., Foote, S.J., Fulton, K. and Whyte, L.G. (2020) Thiomicrorhabdus streamers and sulfur cycling in perennial hypersaline cold springs in the Canadian high Arctic. Environ Microbiol. -Maillard, A., Hübscher, C., Benkhelil, J. and Tahchi, E. (2011) Deformed Messinian markers in the Cyprus Arc: tectonic and/or Messinian Salinity Crisis indicators? Basin Research 23, 146-170. -Mancini, A., Grelaud, M., Ziveri, P., Nallino, E. and Lozar, F. (2021) Calcareous Nannofossil Size and Abundance Response to the Messinian Salinity Crisis Onset and Paleoenvironmental Dynamics. Paleoceanography and Paleoclimatology. -Manzi, V., Argnani, A., Corcagnani, A., Lugli, S. and Roveri, M. (2020) The Messinian salinity crisis in the Adriatic foredeep: Evolution of the largest evaporitic marginal basin in the Mediterranean. Marine and Petroleum Geology 115, 104288.

-Manzi, V., Gennari, R., Hilgen, F., Krijgsman, W., Lugli, S., Roveri, M. and Sierro, F.J. (-Manzi, V., Lugli, S., Roveri, M., Dela Pierre, F., Gennari, R., Lozar, F., Natalicchio, M.,

[START_REF] Manzi | The Messinian salinity crisis in Cyprus: A further step towards a new stratigraphic framework for Eastern Mediterranean[END_REF]

The Messinian salinity crisis in Cyprus:

-

  McKenzie, J., JA, M. and TE, R. (1978) Stable isotopic investigation of carbonate samples related to the Messinian salinity crisis from DSDP Leg 42A, Mediterranean Sea.-[START_REF] Melegy | Microstructure and geochemistry studies on Messinian gypsum deposits from the Northern Coast of Egypt[END_REF] Microstructure and geochemistry studies on Messinian gypsum deposits from the Northern Coast of Egypt. Arabian Journal of Geosciences 7. -Mélières, F., Chamley, H., Coumes, F. and Rouge, P. (1978) X-ray mineralogy studies, Leg 42A, deep sea drilling project, Mediterranean Sea. DSDP 42, 361-383. -Menard., H.W., Smith., S.M. and Pratt., R.M. (1965) The Rhône Deep-Sea Fan in: Whittard., W.F., Bradshaw., R. (Eds.), Submarine gelogy and gophysics. Proceedingsof the Seventeeth Symposium of the Colston Research Society, University of Bristol, pp. 271-285. -Mengdi, Y., Liu, C.-Q., Li, X.-D., Ding, S., Gaoyang, C., Teng, H., Lv, H., Wang, Y., Zhang, X. and Guan, T. (2022) Carbon-sulfur coupling in a seasonally hypoxic, high-sulfate reservoir in SW China: Evidence from stable C S isotopes and sulfate-reducing bacteria. Science of The Total Environment 828, 154537.

	Main text bibliography

___________________________________________________________________________________________ 295 --Micallef, A., Camerlenghi, A., Georgiopoulou, A., Garcia-Castellanos, D., Gutscher, M.-A., Lo Iacono, C., Huvenne, V.A.I., Mountjoy, J.J., Paull, C.K., Le Bas, T., Spatola, D., Facchin, L. and Accettella, D. (2018) Geomorphic evolution of the Malta Escarpment and implications for the Messinian evaporative drawdown in the eastern Mediterranean Sea. Geomorphology 327, 264-283. -Michaud, A.B., Laufer, K., Findlay, A., Pellerin, A., Antler, G., Turchyn, A.V., Røy, H., Wehrmann, L.M. and Jørgensen, B.B. (2020) Glacial influence on the iron and sulfur cycles in Arctic fjord sediments (Svalbard). Geochimica et Cosmochimica Acta 280, 423-440.

  ) Pyrite sulfur isotopes reveal glacial-interglacial environmental changes. Proceedings of the National Academy of Sciences 114, 5941-5945.-Passier, H.F., Middelburg, J.J., de Lange, G.J. and Böttcher, M.E. (1999) Modes of sapropel formation in the eastern Mediterranean: some constraints based on pyrite properties. Marine geology 153, 199-219. Massdependent sulfur isotope fractionation during reoxidative sulfur cycling: A case study from Mangrove Lake, Bermuda. Geochimica et Cosmochimica Acta 149, 152-164. -Pellerin, A., Gilad, A., Holm, S., Findlay, A., Crockford, P., Turchyn, A., Jørgensen, B. and Finster, K. (2019) Large sulfur isotope fractionation by bacterial sulfide oxidation. Science Advances 5. Petsch, S.T. and Berner, R.A. (1998) Coupling the geochemical cycles of C, P, Fe, and S; the effect on atmospheric O <sub>2</sub> and the isotopic records of carbon and sulfur. American Journal of Science 298, 246-262. Picard, A., Gartman, A., Clarke, D.R. and Girguis, P.R. (2018) Sulfate-reducing bacteria influence the nucleation and growth of mackinawite and greigite. Geochimica et Cosmochimica Acta 220, 367-384. -Pierre, C. (1974) Contribution a l'étude sedimentologique et isotopique des evaporites messiniennes de la Mediterranée; implications géodynamiques, Geologie dynamique -Geodynamique sous-marine. Université de Paris VI, p. 118. -Pierre, C. (1982) Teneurs en isotopes stables (1 8 O, 2 H, 1 3C, 3 4S) et conditions de génèse des évaporites marines: application à quelques milieux actuels et au Messinien de la Méditerranée. éditeur inconnu. -Pierre, C. (2018) The isotopic record of gypsum diagenesis in diluted solutions: Observations in natural salinas and experiments. Chemical Geology 493. -Pierre, C. and Fontes, J.-C. (1982) Etude isotopique des saumures et des gypses des marais salants de Salin-de-Giraud (Sud de la France). Géologie Méditerranéenne 9, 479-486. -Pierre, C. and Fontes, J.C. (1978) Isotope composition of Messinian sediments from the Mediterranean Sea as indicators of paleoenvironments and diagenesis.

	-Perrodon, A. (1957) Etude géologique des bassins néogènes sublittoraux de l'Algérie occidentale. -Main text bibliography

-Paytan, A.,

Kastner, M., Campbell, D. and Thiemens, M.H. (1998)

Sulfur Isotopic Composition of Cenozoic Seawater Sulfate. Science 282, 1459-1462. -Pellegrino, L., Dela Pierre, F., Jordan, R.W., Abe, K., Mikami, Y., Natalicchio, M., Gennari, R., Lozar, F. and Carnevale, G. (2020) The upper Miocene diatomaceous sediments of the northernmost Mediterranean region: A lamina-scale investigation of an overlooked palaeoceanographic archive. Sedimentology 67, 3389-3421. -Pellegrino, L., Dela Pierre, F., Natalicchio, M. and Carnevale, G. (2018) The Messinian diatomite deposition in the Mediterranean region and its relationships to the global silica cycle. Earth-Science Reviews 178, 154-176. -Pellegrino, L., Natalicchio, M., Abe, K., Jordan, R., Longo, S., Ferrando, S., Carnevale, G. and Dela Pierre, F. (2021) Tiny, glassy, and rapidly trapped: The nano-sized planktic diatoms in Messinian (late Miocene) gypsum. Geology 49. -Pellerin, A., Antler, G., Røy, H., Findlay, A., Beulig, F., Scholze, C., Turchyn, A.V. and Jørgensen, B.B. (2018) The sulfur cycle below the sulfate-methane transition of marine sediments. Geochimica et Cosmochimica Acta 239, 74-89. -Pellerin, A., Bui, T.H., Rough, M., Mucci, A., Canfield, D.E. and Wing, B.A. (2015) ___________________________________________________________________________________________ 298 --Pierre, C., Purser, B.H., Fontes, J.-C., Busson, G., Tardy, Y., Fritz, P., Lalou, C. and Directeur de thèse, i. (1982) Teneurs en isotopes stables ( ¹ ⁸ O, ² H, ¹ ³C, ³ ⁴S) et conditions de génèse des évaporites marines application à quelques milieux actuels et au Messinien de la Méditerranée, pp. 1 vol. (266-XXIV p.-213 f. de pl.).

  Santarelli, A., Brinkhuis, H., Hilgen, F.J., Lourens, L.J., Versteegh, G.J.M. and Visscher, H. (1998) Orbital signatures in a late Miocene dinoflagellate record from Crete (Greece). Marine Micropaleontology 33, 273-297.-Schaeffer, P., Harrison, W.N., Keely, B.J. and Maxwell, J.R. (1995) Product distributions from chemical degradation of kerogens from a marl from a Miocene evaporitic sequence (Vena del Gesso, N. Italy). Organic Geochemistry 23, 541-554.-Scheiderich, K.,Zerkle, A.L., Helz, G.R., Farquhar, J. and Walker, R.J. (2010) Molybdenum isotope, multiple sulfur isotope, and redox-sensitive element behavior in early Pleistocene Mediterranean sapropels. Chemical Geology 279, 134-144. Schopf, J.W., Farmer, J.D., Foster, I.S., Kudryavtsev, A.B., Gallardo, V.A. and Espinoza, C. (2012) Gypsum-permineralized microfossils and their relevance to the search for life on Mars. Astrobiology 12, 619-633. -Schrader, H. and Gersonde, R. (1978) The late Messinian Mediterranean brackish to freshwater environment diatom floral evidence. In Zachariasse etb al. Microplaeontological counting methods and In Hsü, K., Montadert, L., et al. Init. Repts DSDPPt. 1): Washington (US Govt. Printing Office) 42, 761-775. -Seidenkrantz, M.S., Kouwenhoven, T.J., Jorissen, F.J., Shackleton, N.J. and van der Zwaan, G.J. (2000) Benthic foraminifera as indicators of changing Mediterranean-Atlantic water exchange in the late Miocene. Marine Geology 163, 387-407. -Seitaj, D., Schauer, R., Sulu-Gambari, F., Hidalgo-Martinez, S., Malkin, S.Y., Burdorf, L.D.W., Slomp, C.P. and Meysman, F.J.R. (2015) Cable bacteria generate a firewall against euxinia in seasonally hypoxic basins. Proceedings of the National Academy of Sciences 112, 13278-13283. Sim, M.S., Ono, S., Donovan, K., Templer, S.P. and Bosak, T. (2011b) Effect of electron donors on the fractionation of sulfur isotopes by a marine Desulfovibrio sp. Geochimica et Cosmochimica Acta 75, 4244-4259. -Spear, N., Holland, H.D., Garcia-Veígas, J., Lowenstein, T.K., Giegengack, R. and Peters, H. (2014) Analyses of fluid inclusions in Neoproterozoic marine halite provide oldest measurement of seawater chemistry. Geology 42, 103-106. Tekin, E., Varol, B. and Ayyıldız, T. (2010) Sedimentology and paleoenvironmental evolution of Messinian evaporites in the Iskenderun-Hatay basin complex, Southern Turkey. Sedimentary Geology 229, 282-298. -Thamdrup, B., Finster, K., Fossing, H., Hansen, J.W. and Jørgensen, B.B. (1994) Thiosulfate and sulfite distributions in porewater of marine sediments related to manganese, iron, and sulfur geochemistry. Geochimica et Cosmochimica Acta 58, 67-73. -Thamdrup, B., Finster, K., Hansen, J.W. and Bak, F. (1993) Bacterial Disproportionation of Elemental Sulfur Coupled to Chemical Reduction of Iron or Manganese. Applied and Environmental Microbiology 59, 101-108. -Thode, H.G. and Monster, J. (1965) Sulfur-Isotope Geochemistry of Petroleum, Evaporites, and Ancient Seas1, in: Young, A., Galley, J.E. (Eds.), Fluids in Subsurface Environments. American Association of Petroleum Geologists, p. 0.

	Main text bibliography Main text bibliography
	-Stam, M.C., Mason, P.R.D., Laverman, A., Pallud, C. and Van Cappellen, P. (2006)
	34S/32S fractionation by sulfate-reducing microbial communities in estuarine
	sediments. Geochmica et Cosmochimica Acta 70.

Sansjofre, P., Cartigny, P., Trindade, R.I., Nogueira, A.C., Agrinier, P. and Ader, M. (2016)

Multiple sulfur isotope evidence for massive oceanic sulfate depletion in the aftermath of Snowball Earth. Nature communications 7, 1-8. -___________________________________________________________________________________________ 302 --Selli, R. (1964) Il Messiniano Mayer-Eymar 1867. Proposta di un nostratotype. Giorn. Geol 28, 1-34. -Sghari, A. (2012) Un dépôt évaporitique du Miocène supérieur dans le Sud tunisien : une conséquence de la crise messinienne ? Géomorphologie : relief, processus, environnement 18, 279-294. -Shields, G.A. and Mills, B.J.W. (2020) Evaporite weathering and deposition as a longterm climate forcing mechanism. Geology 49, 299-303. -Sierro, F.J., Flores, J.A., Francés, G., Vazquez, A., Utrilla, R., Zamarreño, I., Erlenkeuser, H. and Barcena, M.A. (2003) Orbitally-controlled oscillations in planktic communities and cyclic changes in western Mediterranean hydrography during the Messinian. Palaeogeography, Palaeoclimatology, Palaeoecology 190, 289-316. -Sigl, W., Chamley, H., Fabricius, F., d'Argoud, G.G. and Müller, J. 13.2 SEDIMENTOLOGY AND ENVIRONMENTAL CONDITIONS OF SAPROPELS. -Sim, M.S., Bosak, T. and Ono, S. (2011a) Large sulfur isotope fractionation does not require disproportionation. Science 333, 74-77. -___________________________________________________________________________________________ 303 --Thode, H.G., Monster, J. and Dunford, H.B. (1961) Sulphur isotope geochemistry. Geochimica et Cosmochimica Acta 25, 159-174. -Thomazo, C., Brayard, A., Elmeknassi, S., Vennin, E., Olivier, N., Caravaca, G., Escarguel, G., Fara, E., Bylund, K.G., Jenks, J.F., Stephen, D.A., Killingsworth, B., Sansjofre, P. and Cartigny, P. (2019) Multiple sulfur isotope signals associated with the late Smithian event and the Smithian/Spathian boundary. Earth-Science Reviews 195, 96-113.

Table E .

 E 1 Oxygen and multiple sulfur isotope composition of the sulfate ion in equilibrium with calcium sulfate minerals from DSDP/ODP sites 134, 374, 376 and 654 and concentrations and multiple sulfur isotopic composition of iron sulfide preserved in the calcium sulfate and marly interbeds from the same cores. Error is 1σ.

	Lithofacies	Sample label	δ 18 OSO4 ±0.2‰	δ 34 SSO4 ±0.2‰	∆ 33 SSO4 ±0.004‰	δ 34 SFeS ±0.4‰	∆ 33 SFeS ±0.008‰	[ FeS ] ±100ppm
	Lithofacies Laminated gypsum dolomitic/pyritic marl-grey	Sample label IODP-654-28A IODP-654-28B	δ 18 OSO4 ±0.2‰ 13,2	δ 34 SSO4 ±0.2‰ 21,7	∆ 33 SSO4 ±0.004‰ 0,026	δ 34 SFeS ±0.4‰ 10,0	∆ 33 SFeS ±0.008‰ 0,054	[ FeS ] 25 ±100ppm
	marl laminated marl-dark grey IODP-654-28C IODP-134-7 with anhydrite anhydrite IODP-134-8 Laminated gypsum IODP-654-29A	13.1	20.5 20,7	0.033 0,024	-33.0 -25,4 -23.0	0.116 0,092 0.112	3335 750 1437
	anhydrite marl marl-brown/yellowish marl-dark grey IODP-654-29C IODP-134-10A IODP-134-10B IODP-654-29B Laminated gypsum IODP-374-17A marl IODP-374-17B Laminated IODP-654-30 gypsum gypsrudite IODP-374-17C marl-brown/yellowish IODP-654-31A	13.5 15,5 13,9 12.3	21.0 21,2 21.2	0.022 0,021 0.036	-11.4 -9.9 -29,4 -21.4 -30.0 -32,9 -5.9	0.102 0.097 0,141 0.097 0.120 0,116 0.043	1714 3306 1001 3549
	porfiroblastic gypsarenite Laminated gypsum	IODP-374-18 IODP-654-31B	10.6 13,1	21.8 20,6	0.041 0,031	4.4 8,5	0.043 0,047	66 24
	dolomitic marl mudstone porfiroblastic gypsarenite Laminated gypsum	IODP-374-19A IODP-654-32A IODP-374-19B IODP-654-32B	11.5 14,5	21.5 20,3	0.031 0,013	-22.0 -38,8 5.9 -28,0	0.048 0,097 0.041 0,048	848 1612 36 251
	laminar anhydrite marl-grey porfiroblastic gypsarenite marl-very dark grey Laminated gypsum Laminated gypsum massive Laminated gypsum anhydrite marl marl-dark grey gypsum silitite IODP-376-16B IODP-374-19C IODP-654-33 IODP-374-20A IODP-654-34A IODP-374-20B IODP-654-34B IODP-374-21 IODP-654-34C IODP-376-16A IODP-654-35 Laminated gypsum IODP-654-36A	10.5 10.8 13,4 9.8 13,7 12.6 10,7	20.7 20.5 20,9 20.7 21,6 21.0 19,3	0.044 0.030 0,019 0.043 0,026 0.029 0,028	-11.7 -40,3 5.5 -38,5 4.1 9,4 -26.4 -26,1 -27.7 -31,2 -15.2 -22,3	0,094 0.103 0.049 0,107 0,041 0.053 0.061 0,096 0,121 0.099 0.025 0,111	1157 6927 74 25 82 9004 46 1043
	gypsrudite gypsum laminated marl-very dark grey	IODP-376-16C IODP-376-16D IODP-654-36B	12.5 13.0	21.5 21.5	0.035 0.027	10.8 6.7 -27,5	0.057 0.054 0,148	21 15
	marl	IODP-376-17A				-21.5	0.090
	gypsum silitite IODP-376-17B	11.6	21.7	0.037	13.2	0.090	16
	gypsrudite	IODP-376-18A	12.4	21.0	0.028	2.6	0.037	10
	gypsum silitite IODP-376-18B	11.6	20.8	0.041	9.0	0.063	19
	gypsrudite	IODP-376-19A	12.6	21.1	0.022	-2.2	0.074	66
	laminated gypsum	IODP-376-19B	15.7	24.4	0.030	-2.8	0.067
	porfiroblastic gypsarenite	IODP-376-20A	12.5	26.2	0.041	-4.9	0.075
	laminated gypsum	IODP-376-20B	9.5	22.3	0.035	-18.2	0.100
	gypsum silitite IODP-376-20C				-17.9	0.086
	entherolithic anhydrite	IODP-376-22	8.6	21.8	0.041	-11.1	0.111

Table B .

 B 1 Sedimentary archives pyrite dataset and references used in Chapter 4 Table B. 2 Surifical modern sediment dataset and references used in Chapter 4

	Water Depth (mbsl) 2192 2192 2192 2192 2192 2192 2192 2192 2192 2192 2192 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 Water Depth (mbsl) 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 Water Depth (mbsl) 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 Water Depth (mbsl) 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 Water Depth (mbsl) 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 298,5 Water Depth (mbsl) 298,5 298,5 298,5 298,5 298,5 298,5 298,5 348 348 348 348 348 348 348 348 348 348 348 348 348 348 348 348 Water Depth (mbsl) Water Depth (mbsl) Water Depth (mbsl) Water Depth (mbsl) 348 348 348 348 348 348 348 348 3290 3290 3290 3290 3290 3290 3290 3290 3290 3290 3290 3290 3290 Water Depth (mbsl) Water Depth (mbsl) Water Depth (mbsl) Water Depth (mbsl) 150 150 150 Water Depth (mbsl) Water Depth (mbsl) Water Depth (mbsl) Water Depth (mbsl) Water Depth (mbsl) Water Depth (mbsl) Water Depth (mbsl) Water Depth (mbsl) Water Depth (mbsl) Water Depth (mbsl) Water Depth (mbsl) Water Depth (mbsl) Water Depth (mbsl) Water Depth (mbsl) Water Depth (mbsl) Water Depth (mbsl) Water Depth (mbsl) Water Depth (mbsl) Water Depth (mbsl) Water Depth (mbsl) Water Depth (mbsl) 1760 27.15 -44,94 Lat. (°) 33,84 33,84 33,84 33,84 33,84 33,84 33,84 33,84 33,84 33,84 33,84 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 Lat. (°) 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 Lat. (°) 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 Lat. (°) 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 Lat. (°) 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 42,69 Lat. (°) 42,69 42,69 42,69 42,69 42,69 42,69 42,69 -44,94 -44,94 -44,94 -44,94 -44,94 -44,94 -44,94 -44,94 -44,94 -44,94 -44,94 -44,94 -44,94 -44,94 -44,94 Lat. (°) -44,94 -44,94 -44,94 -44,94 -44,94 -44,94 -44,94 -44,94 -44,94 -44,94 -44,94 -44,94 -44,94 -44,94 -44,94 -44,94 -44,94 -44,94 -44,94 -44,94 -44,94 Lat. (°) -44,94 -44,94 -44,94 -44,94 -44,94 -44,94 -44,94 -44,94 -44,94 -44,94 -44,94 -44,94 -44,94 -44,94 -44,94 -44,94 -44,94 -44,94 -44,94 -44,94 -44,94 Lat. (°) -44,94 -44,94 -44,94 -44,94 -44,94 -44,94 -44,94 -44,94 -44,94 -44,94 -44,94 -44,94 -44,94 -44,94 -44,94 -44,94 -44,94 -44,94 -44,94 -44,94 -44,94 Lat. (°) -44,94 -44,94 -44,94 -44,94 -44,94 -44,94 -44,94 -44,94 -41,79 -41,79 -41,79 -41,79 -41,79 -41,79 -41,79 -41,79 -41,79 -41,79 -41,79 -41,79 -41,79 Lat. (°) -41,79 -41,79 -41,79 -41,79 -41,79 -41,79 -41,79 -41,79 -41,79 -41,79 -41,79 -41,79 -41,79 -41,79 -41,79 -41,79 -41,79 -41,79 -41,79 -41,79 -41,79 Lat. (°) -41,79 -41,79 -41,79 -41,79 -41,79 -41,79 -41,79 -41,79 -41,79 -41,79 -41,79 -41,79 -41,79 -41,79 -41,79 -41,79 -41,79 -41,79 -41,79 -41,79 -41,79 Lat. (°) -41,79 -41,79 -41,79 -41,79 -41,79 -41,79 -41,79 -41,79 -41,79 -41,79 -41,79 -41,79 -41,79 -41,79 -41,79 -41,79 -41,79 -41,79 -41,79 -41,79 -41,79 Lat. (°) -41,79 -41,79 -41,79 -41,79 -41,79 -41,79 -41,79 -41,79 -41,79 -41,79 -41,79 -41,79 -41,79 -41,79 34,07 34,07 34,07 33,84 33,84 33,84 -10,98 -10,98 -10,98 Lat. (°) -10,98 -10,98 -10,98 -10,98 -10,98 -10,98 -10,98 -10,98 -10,98 -10,98 -10,98 -10,98 -10,98 -10,98 -10,98 -10,98 -10,98 -10,98 -10,98 -10,98 -10,98 -10,98 -10,98 -10,98 -10,98 -10,98 -10,98 Lat. (°) -10,98 -10,98 -10,98 -10,98 -10,98 -10,98 -10,98 -10,98 -10,98 -10,98 -10,98 -10,98 -10,98 -10,98 -10,98 -10,98 -10,98 -10,98 -10,98 -10,98 -10,98 -10,98 -10,98 -10,98 -10,98 -10,98 -10,98 Lat. (°) -10,98 -10,98 -10,98 -10,98 -10,98 -10,98 16,03 16,03 16,03 16,03 16,03 16,03 16,03 16,03 16,03 16,03 16,03 16,03 16,03 16,03 16,03 16,03 16,03 16,03 16,03 16,03 16,03 Lat. (°) 16,03 16,03 16,03 16,03 16,03 16,03 16,03 16,03 16,03 16,03 16,03 16,03 16,03 16,03 16,03 16,03 16,03 16,03 16,03 16,03 16,03 16,03 16,03 16,03 16,03 16,03 16,03 Lat. (°) 16,03 16,03 16,03 16,03 16,03 16,03 16,03 16,03 16,03 16,03 16,03 16,03 16,03 16,03 16,03 16,03 16,03 16,03 16,03 Lat. (°) Long. (°) Lat. (°) Long. (°) Lat. (°) Long. (°) Lat. (°) Long. (°) Lat. (°) Long. (°) Lat. (°) Long. (°) Lat. (°) Long. (°) Lat. (°) Long. (°) Lat. (°) Long. (°) Lat. (°) Long. (°) Lat. (°) Long. (°) Lat. (°) Long. (°) Lat. (°) Long. (°) Lat. (°) Long. (°) Lat. (°) Long. (°) Lat. (°) Long. (°) 43.85 30.17 -110.87 1760 27.15 -110.87 1170 25.53 -110.55 3260 26.35 -109.93 2860 23.78 -108.67 23.38 -106.95 1000 23.23 -107.12 43.71 30.10 43.71 30.10 43.71 30.10 43.71 30.10 43.71 30.10 43.67 30.13 43.67 30.13 43.67 30.13 43.67 30.13 43.67 30.13 1176 43.53 30.22 1176 43.53 30.22 1176 43.53 30.22 1176 43.53 30.22 -10.03 -79.07 -10.03 -79.07 -10.03 -79.07 -10.03 -79.07 3821 -38.22 -53.35 0.40 Long. (°) 24,88 24,88 24,88 24,88 24,88 24,88 24,88 24,88 24,88 24,88 24,88 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 Long. (°) 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 Long. (°) 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 Long. (°) 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 Long. (°) 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 3,84 Long. (°) 3,84 3,84 3,84 3,84 3,84 3,84 3,84 172,02 172,02 172,02 172,02 172,02 172,02 172,02 172,02 172,02 172,02 172,02 172,02 172,02 172,02 172,02 172,02 Long. (°) 172,02 172,02 172,02 172,02 172,02 172,02 172,02 172,02 172,02 172,02 172,02 172,02 172,02 172,02 172,02 172,02 172,02 172,02 172,02 172,02 172,02 Long. (°) 172,02 172,02 172,02 172,02 172,02 172,02 172,02 172,02 172,02 172,02 172,02 172,02 172,02 172,02 172,02 172,02 172,02 172,02 172,02 172,02 172,02 Long. (°) 172,02 172,02 172,02 172,02 172,02 172,02 172,02 172,02 172,02 172,02 172,02 172,02 172,02 172,02 172,02 172,02 172,02 172,02 172,02 172,02 172,02 Long. (°) 172,02 172,02 172,02 172,02 172,02 172,02 172,02 172,02 -171,50 -171,50 -171,50 -171,50 -171,50 -171,50 -171,50 -171,50 -171,50 -171,50 -171,50 -171,50 -171,50 Long. (°) -171,50 -171,50 -171,50 -171,50 -171,50 -171,50 -171,50 -171,50 -171,50 -171,50 -171,50 -171,50 -171,50 -171,50 -171,50 -171,50 -171,50 -171,50 -171,50 -171,50 -171,50 Long. (°) -171,50 -171,50 -171,50 -171,50 -171,50 -171,50 -171,50 -171,50 -171,50 -171,50 -171,50 -171,50 -171,50 -171,50 -171,50 -171,50 -171,50 -171,50 -171,50 -171,50 -171,50 Long. (°) -171,50 -171,50 -171,50 -171,50 -171,50 -171,50 -171,50 -171,50 -171,50 -171,50 -171,50 -171,50 -171,50 -171,50 -171,50 -171,50 -171,50 -171,50 -171,50 -171,50 -171,50 Long. (°) -171,50 -171,50 -171,50 -171,50 -171,50 -171,50 -171,50 -171,50 -171,50 -171,50 -171,50 -171,50 -171,50 -171,50 32,73 Passier et al. (1999) Reference Scheiderich et al. (2010) Scheiderich et al. (2010) Scheiderich et al. (2010) Scheiderich et al. (2010) Scheiderich et al. (2010) Scheiderich et al. (2010) Scheiderich et al. (2010) Scheiderich et al. (2010) Scheiderich et al. (2010) Scheiderich et al. (2010) Scheiderich et al. (2010) Pasquier et al. (2017) Pasquier et al. (2017) Pasquier et al. (2017) Pasquier et al. (2017) Pasquier et al. (2017) Pasquier et al. (2017) Pasquier et al. (2017) Pasquier et al. (2017) Pasquier et al. (2017) Pasquier et al. (2017) Pasquier et al. (2017) Pasquier et al. (2017) Reference Pasquier et al. (2017) Pasquier et al. (2017) Pasquier et al. (2017) Pasquier et al. (2017) Pasquier et al. (2017) Pasquier et al. (2017) Pasquier et al. (2017) Pasquier et al. (2017) Pasquier et al. (2017) Pasquier et al. (2017) Pasquier et al. (2017) Pasquier et al. (2017) Pasquier et al. (2017) Pasquier et al. (2017) Pasquier et al. (2017) Pasquier et al. (2017) Pasquier et al. (2017) Pasquier et al. (2017) Pasquier et al. (2017) Pasquier et al. (2017) Pasquier et al. (2017) Pasquier et al. (2017) Pasquier et al. (2017) Pasquier et al. (2017) Pasquier et al. (2017) Pasquier et al. (2017) Pasquier et al. (2017) Reference Pasquier et al. (2017) Pasquier et al. (2017) Pasquier et al. (2017) Pasquier et al. (2017) Pasquier et al. (2017) Pasquier et al. (2017) Pasquier et al. (2017) Pasquier et al. (2017) Pasquier et al. (2017) Pasquier et al. (2017) Pasquier et al. (2017) Pasquier et al. (2017) Pasquier et al. (2017) Pasquier et al. (2017) Pasquier et al. (2017) Pasquier et al. (2017) Pasquier et al. (2017) Pasquier et al. (2017) Pasquier et al. (2017) Pasquier et al. 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(2021a) Pasquier et al. (2021a) 32,73 Passier et al. (1999) 32,73 Passier et al. (1999) 24,88 Passier et al. (1999) 24,88 Passier et al. (1999) 24,88 Passier et al. (1999) -77,96 Meister et al. (2019) -77,96 Meister et al. (2019) -77,96 Meister et al. (2019) Long. (°) Reference -77,96 Meister et al. (2019) -77,96 Meister et al. (2019) -77,96 Meister et al. (2019) -77,96 Meister et al. (2019) -77,96 Meister et al. (2019) -77,96 Meister et al. (2019) -77,96 Meister et al. (2019) -77,96 Meister et al. (2019) -77,96 Meister et al. (2019) -77,96 Meister et al. (2019) -77,96 Meister et al. (2019) -77,96 Meister et al. (2019) -77,96 Meister et al. (2019) -77,96 Meister et al. (2019) -77,96 Meister et al. (2019) -77,96 Meister et al. (2019) -77,96 Meister et al. (2019) -77,96 Meister et al. (2019) -77,96 Meister et al. (2019) -77,96 Meister et al. (2019) -77,96 Meister et al. (2019) -77,96 Meister et al. (2019) -77,96 Meister et al. (2019) -77,96 Pasquier et al. (2021b) -77,96 Pasquier et al. (2021b) -77,96 Pasquier et al. (2021b) -77,96 Pasquier et al. (2021b) Long. (°) Reference -77,96 Pasquier et al. (2021b) -77,96 Pasquier et al. (2021b) -77,96 Pasquier et al. (2021b) -77,96 Pasquier et al. (2021b) -77,96 Pasquier et al. (2021b) -77,96 Pasquier et al. (2021b) -77,96 Pasquier et al. (2021b) -77,96 Pasquier et al. (2021b) -77,96 Pasquier et al. (2021b) -77,96 Pasquier et al. (2021b) -77,96 Pasquier et al. (2021b) -77,96 Pasquier et al. (2021b) -77,96 Pasquier et al. (2021b) -77,96 Pasquier et al. (2021b) -77,96 Pasquier et al. (2021b) -77,96 Pasquier et al. (2021b) -77,96 Pasquier et al. (2021b) -77,96 Pasquier et al. (2021b) -77,96 Pasquier et al. (2021b) -77,96 Pasquier et al. (2021b) -77,96 Pasquier et al. (2021b) -77,96 Pasquier et al. (2021b) -77,96 Pasquier et al. (2021b) -77,96 Pasquier et al. (2021b) -77,96 Pasquier et al. (2021b) -77,96 Pasquier et al. (2021b) -77,96 Pasquier et al. (2021b) Long. (°) Reference -77,96 Pasquier et al. (2021b) -77,96 Pasquier et al. (2021b) -77,96 Pasquier et al. (2021b) -77,96 Pasquier et al. (2021b) -77,96 Pasquier et al. (2021b) -77,96 Pasquier et al. (2021b) 81,71 Peketi et al. (2015) 81,71 Peketi et al. (2015) 81,71 Peketi et al. (2015) 81,71 Peketi et al. (2015) 81,71 Peketi et al. (2015) 81,71 Peketi et al. (2015) 81,71 Peketi et al. (2015) 81,71 Peketi et al. (2015) 81,71 Peketi et al. (2015) 81,71 Peketi et al. (2015) 81,71 Peketi et al. (2015) 81,71 Peketi et al. (2015) 81,71 Peketi et al. (2015) 81,71 Peketi et al. (2015) 81,71 Peketi et al. (2015) 81,71 Peketi et al. (2015) 81,71 Peketi et al. (2015) 81,71 Peketi et al. (2015) 81,71 Peketi et al. (2015) 81,71 Peketi et al. (2015) 81,71 Peketi et al. (2015) Long. (°) Reference 81,71 Peketi et al. (2015) 81,71 Peketi et al. (2015) 81,71 Peketi et al. (2015) 81,71 Peketi et al. (2015) 81,71 Peketi et al. (2015) 81,71 Peketi et al. (2015) 81,71 Peketi et al. (2015) 81,71 Peketi et al. (2015) 81,71 Peketi et al. (2015) 81,71 Peketi et al. (2015) 81,71 Peketi et al. (2015) 81,71 Peketi et al. (2015) 81,71 Peketi et al. (2015) 81,71 Peketi et al. (2015) 81,71 Peketi et al. (2015) 81,71 Peketi et al. (2015) 81,71 Peketi et al. (2015) 81,71 Peketi et al. (2015) 81,71 Peketi et al. (2015) 81,71 Peketi et al. (2015) 81,71 Peketi et al. (2015) 81,71 Peketi et al. (2015) 81,71 Peketi et al. (2015) 81,71 Peketi et al. (2015) 81,71 Peketi et al. (2015) 81,71 Peketi et al. (2015) 81,71 Peketi et al. (2015) Long. (°) Reference 81,71 Peketi et al. (2015) 81,71 Peketi et al. (2015) 81,71 Peketi et al. (2015) 81,71 Peketi et al. (2015) 81,71 Peketi et al. (2015) 81,71 Peketi et al. (2015) 81,71 Peketi et al. (2015) 81,71 Peketi et al. (2015) 81,71 Peketi et al. (2015) 81,71 Peketi et al. (2015) 81,71 Peketi et al. (2015) 81,71 Peketi et al. (2015) 81,71 Peketi et al. (2015) 81,71 Peketi et al. (2015) 81,71 Peketi et al. (2015) 81,71 Peketi et al. (2015) 81,71 Peketi et al. (2015) 81,71 Peketi et al. (2015) 81,71 Peketi et al. (2015) Depth (mbsf) Reference Depth (mbsf) Reference Depth (mbsf) Reference Depth (mbsf) Reference Depth (mbsf) Reference Depth (mbsf) Reference Depth (mbsf) Reference Depth (mbsf) Reference Depth (mbsf) Reference Depth (mbsf) Reference Depth (mbsf) Reference Depth (mbsf) Reference Depth (mbsf) Reference Depth (mbsf) Reference Depth (mbsf) Reference Depth (mbsf) Reference 1.00 Jørgensen (2005) 0.23 Lein et al. (1976b) 0.55 Lein et al. (1976b) 0.63 Lein et al. (1976b) 0.43 Lein et al. (1976b) 0.40 Lein et al. (1976b) 0.70 Lein et al. (1976b) 0.90 Lein et al. (1976b) 0.00 Lein et al. (1976b) 0.40 Lein et al. (1976b) 0.60 Lein et al. (1976b) 0.80 Lein et al. (1976b) 1.00 Lein et al. (1976b) 0.20 Lein et al. (1976b) 0.40 Lein et al. (1976b) 0.60 Lein et al. (1976b) 0.80 Lein et al. (1976b) 1.00 Lein et al. (1976b) 0.30 Lein et al. (1976b) 0.50 Lein et al. (1976b) 0.80 Lein et al. (1976b) 1.00 Lein et al. (1976b) Arning et al. (2009) Arning et al. (2009) Arning et al. (2009) Arning et al. (2009) Riedinger et al. (2017) 3821 -38.22 -53.35 0.70 Riedinger et al. (2017) 3821 -38.22 -53.35 0.90 Riedinger et al. (2017) 787 50.52 155.30 0.03 Strizhov et al. (1990) 787 50.52 155.30 0.11 Strizhov et al. (1990) 787 50.52 155.30 0.20 Strizhov et al. (1990) 787 50.52 155.30 0.30 Strizhov et al. (1990) 787 50.52 155.30 0.39 Strizhov et al. (1990) 787 50.52 155.30 0.58 Strizhov et al. (1990) 787 50.52 155.30 1.03 Strizhov et al. (1990) 892.9 10.71 -65.17 0.05 Lyons et al. (2003) 892.9 10.71 -65.17 0.10 Lyons et al. (2003) 892.9 10.71 -65.17 0.15 Lyons et al. (2003) 892.9 10.71 -65.17 0.20 Lyons et al. (2003) 892.9 10.71 -65.17 0.25 Lyons et al. (2003) 892.9 10.71 -65.17 0.30 Lyons et al. (2003) 892.9 10.71 -65.17 0.35 Lyons et al. (2003) 892.9 10.71 -65.17 0.40 Lyons et al. (2003) 892.9 10.71 -65.17 0.55 Lyons et al. (2003) 892.9 10.71 -65.17 0.75 Lyons et al. (2003) 892.9 10.71 -65.17 0.95 Lyons et al. (2003) 892.9 10.71 -65.17 1.05 Lyons et al. (2003) 273 -11.07 -78.08 0.30 Mossmann et al. (1991) 273 -11.07 -78.08 0.80 Mossmann et al. (1991) 458 -13.48 -76.89 0.30 Mossmann et al. (1991) 458 -13.48 -76.89 0.90 Mossmann et al. (1991) 62 -17.22 11.69 0.02 Lein and Ivanov (2009) 62 -17.22 11.69 0.14 Lein and Ivanov (2009) 62 -17.22 11.69 0.38 Lein and Ivanov (2009) 62 -17.22 11.69 0.78 Lein and Ivanov (2009) 62 -17.22 11.69 1.28 Lein and Ivanov (2009) 123 -17.17 11.54 0.02 Lein and Ivanov (2009) 123 -17.17 11.54 0.10 Lein and Ivanov (2009) 965 -17.21 11.17 0.02 Lein and Ivanov (2009) 965 -17.21 11.17 0.10 Lein and Ivanov (2009) 4180 -17.10 9.69 0.01 Lein and Ivanov (2009) 4180 -17.10 9.69 0.10 Lein and Ivanov (2009) 4180 -17.10 9.69 0.60 Lein and Ivanov (2009) 4180 -17.10 9.69 0.75 Lein and Ivanov (2009) 4180 -17.10 9.69 0.98 Lein and Ivanov (2009) 42 -22.89 14.39 0.02 Lein and Ivanov (2009) 42 -22.89 14.39 0.06 Lein and Ivanov (2009) 42 -22.89 14.39 0.15 Lein and Ivanov (2009) 98 -22.94 14.27 0.02 Lein and Ivanov (2009) 98 -22.94 14.27 0.09 Lein and Ivanov (2009) 98 -22.94 14.27 0.25 Lein and Ivanov (2009) 98 -22.94 14.27 0.45 Lein and Ivanov (2009) 98 -22.94 14.27 0.95 Lein and Ivanov (2009) 60 -24.97 14.71 0.04 Lein and Ivanov (2009) 98 -22.94 14.27 1.25 Lein and Ivanov (2009) 60 -24.97 14.71 0.08 Lein and Ivanov (2009) 60 -24.97 14.71 0.13 Lein and Ivanov (2009) 103 -25.01 14.54 0.01 Lein and Ivanov (2009) 103 -25.01 14.54 0.05 Lein and Ivanov (2009) 160 -25.07 14.29 0.05 Lein and Ivanov (2009) 7 45.38 37.42 0.08 Liu et al. (1996) 7 45.47 29.82 0.08 Liu et al. (1996) 2 45.46 29.81 0.03 Liu et al. (1996) 15 45.47 29.83 0.03 Liu et al. (1996) 5 45.39 29.78 0.03 Liu et al. (1996) 5 45.31 29.77 0.03 Liu et al. (1996) 5 46.07 30.48 0.02 Liu et al. (1996) 15 46.00 30.50 0.02 Liu et al. (1996) 11 46.02 30.48 0.02 Liu et al. (1996) 19 54.48 10.75 0.01 Hartmann and Nielsen (1968) 19 54.48 10.75 0.02 Hartmann and Nielsen (1968) 19 54.48 10.75 0.04 Hartmann and Nielsen (1968) 19 54.48 10.75 0.06 Hartmann and Nielsen (1968) 19 54.48 10.75 0.08 Hartmann and Nielsen (1968) 19 54.48 10.75 0.11 Hartmann and Nielsen (1968) 19 54.48 10.75 0.14 Hartmann and Nielsen (1968) 19 54.48 10.75 0.17 Hartmann and Nielsen (1968) 54.48 10.75 0.23 Hartmann and Nielsen (1968) 54.48 10.75 0.28 Hartmann and Nielsen (1968) 54.48 10.75 0.33 Hartmann and Nielsen (1968) 54.48 10.75 0.38 Hartmann and Nielsen (1968) 54.52 10.05 0.01 Hartmann and Nielsen (1968) 54.52 10.05 0.02 Hartmann and Nielsen (1968) 54.52 10.05 0.04 Hartmann and Nielsen (1968) 54.52 10.05 0.06 Hartmann and Nielsen (1968) 54.52 10.05 0.08 Hartmann and Nielsen (1968) 54.52 10.05 0.11 Hartmann and Nielsen (1968) 54.52 10.05 0.14 Hartmann and Nielsen (1968) 54.52 10.05 0.16 Hartmann and Nielsen (1968) 54.52 10.05 0.18 Hartmann and Nielsen (1968) 54.52 10.05 0.22 Hartmann and Nielsen (1968) 54.52 10.05 0.25 Hartmann and Nielsen (1968) 54.52 10.05 0.28 Hartmann and Nielsen (1968) 54.52 10.05 0.31 Hartmann and Nielsen (1968) 54.52 10.05 0.34 Hartmann and Nielsen (1968) 54.52 10.05 0.37 Hartmann and Nielsen (1968) 54.52 10.05 0.40 Hartmann and Nielsen (1968) 54.53 10.05 0.01 Hartmann and Nielsen (1968) 30.5 54.59 10.05 1.05 Hartmann and Nielsen (1968) 55.38 10.05 0.03 Liu et al. (2020) 55.38 10.05 0.08 Liu et al. (2020) 55.38 10.05 0.13 Liu et al. (2020) 55.38 10.05 0.18 Liu et al. (2020) 55.38 10.05 0.23 Liu et al. (2020) 55.38 10.05 0.28 Liu et al. (2020) 55.38 10.05 0.33 Liu et al. (2020) 55.38 10.05 0.40 Liu et al. (2020) 55.38 10.05 0.50 Liu et al. (2020) 55.38 10.05 0.60 Liu et al. (2020) 55.38 10.05 0.70 Liu et al. (2020) 55.38 10.05 0.80 Liu et al. (2020) 55.39 10.05 0.01 Liu et al. (2020) 55.39 10.05 0.03 Liu et al. (2020) 55.39 10.05 0.05 Liu et al. (2020) 55.39 10.05 0.07 Liu et al. (2020) 55.39 10.05 0.09 Liu et al. (2020) 55.39 10.05 0.13 Liu et al. (2020) 55.39 10.05 0.18 Liu et al. (2020) 55.39 10.05 0.23 Liu et al. (2020) 55.39 10.05 0.28 Liu et al. (2020) 55.39 10.05 0.33 Liu et al. (2020) 55.39 10.05 0.38 Liu et al. (2020) 55.39 10.05 0.43 Liu et al. (2020) 55.39 10.05 0.50 Liu et al. (2020) 55.39 10.05 0.60 Liu et al. (2020) 55.39 10.05 0.40 Liu et al. (2020) 55.39 10.05 0.62 Liu et al. (2020) 55.39 10.05 0.80 Liu et al. (2020) 55.39 10.05 0.99 Liu et al. (2020) 55.41 15.48 0.01 Liu et al. (2020) 55.41 15.48 0.03 Liu et al. (2020) 55.41 15.48 0.06 Liu et al. (2020) 55.41 15.48 0.09 Liu et al. (2020) 55.41 15.48 0.13 Liu et al. (2020) 55.41 15.48 0.18 Liu et al. (2020) 55.41 15.48 0.23 Liu et al. (2020) 55.41 15.48 0.28 Liu et al. (2020) 55.41 15.48 0.38 Liu et al. (2020) 55.41 15.48 0.48 Liu et al. (2020) 55.41 15.48 0.58 Liu et al. (2020) 55.41 15.48 0.68 Liu et al. (2020) 55.41 15.48 0.10 Liu et al. (2020) 55.41 15.48 0.23 Liu et al. (2020) 55.41 15.48 0.33 Liu et al. (2020) 55.41 15.48 0.43 Liu et al. (2020) 55.41 15.48 0.50 Liu et al. (2020) 55.41 15.48 0.60 Liu et al. (2020) 55.41 15.48 0.70 Liu et al. (2020) 55.41 15.48 0.80 Liu et al. (2020) 55.41 15.48 0.90 Liu et al. (2020) 55.41 15.48 1.00 Liu et al. (2020) 55.41 15.48 0.43 Liu et al. (2020) 55.41 15.48 0.57 Liu et al. (2020) 55.41 15.48 0.80 Liu et al. (2020) 55.41 15.48 1.00 Liu et al. (2020) 590 34.22 -120.03 0.05 Raven et al. (2016) 590 34.22 -120.03 0.01 Raven et al. (2016) 590 34.22 -120.03 0.04 Raven et al. (2016) 34.22 -120.03 0.11 Raven et al. (2016) 34.22 -120.03 0.16 Raven et al. (2016) 34.22 -120.03 0.21 Raven et al. (2016) 34.22 -120.03 0.26 Raven et al. (2016) 34.22 -120.03 0.31 Raven et al. (2016) 34.22 -120.03 0.36 Raven et al. (2016) 34.22 -120.03 0.41 Raven et al. (2016) 34.22 -120.03 0.46 Raven et al. (2016) 34.22 -120.03 0.51 Raven et al. (2016) 34.22 -120.03 0.56 Raven et al. (2016) 34.22 -120.03 0.61 Raven et al. (2016) 34.22 -120.03 0.66 Raven et al. (2016) 34.22 -120.03 0.23 Raven et al. (2016) 34.22 -120.03 0.43 Raven et al. (2016) 34.22 -120.03 0.50 Raven et al. (2016) 34.22 -120.03 0.55 Raven et al. (2016) 34.22 -120.03 0.58 Raven et al. (2016) 34.22 -120.03 0.63 Raven et al. (2016) 34.22 -120.03 0.73 Raven et al. (2016) 34.22 -120.03 0.78 Raven et al. (2016) 34.22 -120.03 0.85 Raven et al. (2016) 590 34.22 -120.03 0.90 Raven et al. (2016) 590 34.22 -120.03 0.95 Raven et al. (2016) 590 34.22 -120.03 1.00 Raven et al. (2016) 824 29.42 34.91 0.31 Boyko et al. (2021) 824 29.42 34.91 0.43 Boyko et al. (2021) 26.21 35.36 0.01 Boyko et al. (2021) 26.21 35.36 0.03 Boyko et al. (2021) 26.21 35.36 0.05 Boyko et al. (2021) 26.21 35.36 0.07 Boyko et al. (2021) 26.21 35.36 0.09 Boyko et al. (2021) 26.21 35.36 0.11 Boyko et al. (2021) 26.21 35.36 0.13 Boyko et al. (2021) 26.21 35.36 0.15 Boyko et al. (2021) 26.21 35.36 0.17 Boyko et al. (2021) 26.21 35.36 0.19 Boyko et al. (2021) 26.21 35.36 0.21 Boyko et al. (2021) 26.21 35.36 0.23 Boyko et al. (2021) 26.21 35.36 0.25 Boyko et al. (2021) 26.21 35.36 0.27 Boyko et al. (2021) 26.21 35.36 0.29 Boyko et al. (2021) 26.21 35.36 0.31 Boyko et al. (2021) 26.21 35.36 0.33 Boyko et al. (2021) 26.21 35.36 0.35 Boyko et al. (2021) 26.21 35.36 0.37 Boyko et al. (2021) 10.63 93.00 0.21 Liu et al. (2022) 21.91 118.82 0.18 Liu et al. (2022) 21.91 118.82 0.63 Liu et al. (2022) 21.91 118.82 1.07 Liu et al. (2022) 96 55.65 15.78 0.40 Liu et al. (2022) 96 55.65 15.78 0.62 Liu et al. (2022) 55.65 15.78 0.80 Liu et al. (2022) 55.65 15.78 0.99 Liu et al. (2022) 55.65 15.78 1.18 Liu et al. (2022) 55.41 15.48 0.43 Liu et al. (2022) 55.41 15.48 0.57 Liu et al. (2022) 55.41 15.48 0.80 Liu et al. (2022) 55.41 15.48 1.00 Liu et al. (2022) 55.41 15.48 1.17 Liu et al. (2022) 56.11 10.46 0.01 Liu et al. (2022) 56.11 10.46 0.04 Liu et al. (2022) 56.11 10.46 0.08 Liu et al. (2022) 56.11 10.46 0.12 Liu et al. (2022) 56.11 10.46 0.16 Liu et al. (2022) 56.11 10.46 0.20 Liu et al. (2022) 56.11 10.46 0.24 Liu et al. (2022) 56.11 10.46 0.32 Liu et al. (2022) 56.11 10.46 0.36 Liu et al. (2022) 56.11 10.46 0.40 Liu et al. (2022) 56.11 10.46 0.44 Liu et al. (2022) 56.11 10.46 0.48 Liu et al. (2022) 56.11 10.46 0.52 Liu et al. (2022) 56.11 10.46 0.56 Liu et al. (2022) 56.11 10.46 0.60 Liu et al. (2022) 54.52 10.02 0.01 Hartmann and Nielsen (2012) 54.52 10.02 0.02 Hartmann and Nielsen (2012) 54.52 10.02 0.04 Hartmann and Nielsen (2012) 54.52 10.02 0.06 Hartmann and Nielsen (2012) 54.52 10.02 0.12 Hartmann and Nielsen (2012) 54.52 10.02 0.18 Hartmann and Nielsen (2012) 54.52 10.02 0.24 Hartmann and Nielsen (2012) 54.52 10.02 0.30 Hartmann and Nielsen (2012) 54.52 10.02 0.01 Hartmann and Nielsen (2012) 54.52 10.02 0.02 Hartmann and Nielsen (2012) 54.52 10.02 0.04 Hartmann and Nielsen (2012) 54.52 10.02 0.06 Hartmann and Nielsen (2012) 54.52 10.02 0.08 Hartmann and Nielsen (2012) 54.52 10.02 0.11 Hartmann and Nielsen (2012) 54.52 10.02 0.14 Hartmann and Nielsen (2012) 54.52 10.02 0.16 Hartmann and Nielsen (2012) 54.52 10.02 0.18 Hartmann and Nielsen (2012) 54.52 10.02 0.22 Hartmann and Nielsen (2012) 54.52 10.02 0.25 Hartmann and Nielsen (2012) 54.52 10.02 0.28 Hartmann and Nielsen (2012) 54.52 10.02 0.31 Hartmann and Nielsen (2012) 54.52 10.02 0.34 Hartmann and Nielsen (2012) 54.52 10.02 0.37 Hartmann and Nielsen (2012) 54.52 10.02 0.40 Hartmann and Nielsen (2012) 54.52 10.02 0.01 Hartmann and 54.52 10.02 0.02 Hartmann and Nielsen (2012) 110 57.50 19.92 0.04 Lein (1983) 110 57.50 19.92 0.11 Lein (1983) 226 28.63 -112.33 0.80 Sweeney and Kaplan 4974 28.23 -75.00 0.40 Scharff (1980) (1980) San Pedro Martir Basin Locality Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Locality Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Locality Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Locality Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Locality Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Locality Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Locality Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Locality Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Locality Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Locality Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Locality Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Locality Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Locality Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Locality Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Eastern New Zealand margin Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Mediterranean sea Peru Margin Peru Margin Peru Margin Locality Peru Margin Peru Margin Peru Margin Peru Margin Peru Margin Peru Margin Peru Margin Peru Margin Peru Margin Peru Margin Peru Margin Peru Margin Peru Margin Peru Margin Peru Margin Peru Margin Peru Margin Peru Margin Peru Margin Peru Margin Peru Margin Peru Margin Peru Margin Peru Margin Peru Margin Peru Margin Peru Margin Locality Peru Margin Peru Margin Peru Margin Peru Margin Peru Margin Peru Margin Peru Margin Peru Margin Peru Margin Peru Margin Peru Margin Peru Margin Peru Margin Peru Margin Peru Margin Peru Margin Peru Margin Peru Margin Peru Margin Peru Margin Peru Margin Peru Margin Peru Margin Peru Margin Peru Margin Peru Margin Peru Margin Locality Peru Margin Peru Margin Peru Margin Peru Margin Peru Margin Peru Margin Bengal bay Bengal bay Bengal bay Bengal bay Bengal bay Bengal bay Bengal bay Bengal bay Bengal bay Bengal bay Bengal bay Bengal bay Bengal bay Bengal bay Bengal bay Bengal bay Bengal bay Bengal bay Bengal bay Bengal bay Bengal bay Locality Bengal bay Bengal bay Bengal bay Bengal bay Bengal bay Bengal bay Bengal bay Bengal bay Bengal bay Bengal bay Bengal bay Bengal bay Bengal bay Bengal bay Bengal bay Bengal bay Bengal bay Bengal bay Bengal bay Bengal bay Bengal bay Bengal bay Bengal bay Bengal bay Bengal bay Bengal bay Bengal bay Locality Bengal bay Bengal bay Bengal bay Bengal bay Bengal bay Bengal bay Bengal bay Bengal bay Bengal bay Bengal bay Bengal bay Bengal bay Bengal bay Bengal bay Bengal bay Bengal bay Bengal bay Bengal bay Bengal bay Locality Locality Locality Locality Locality Locality Locality Locality Locality Locality Locality Locality Locality Locality Locality Locality Western Black Sea Golfe de Californie Golfe de Californie Golfe de Californie Golfe de Californie Golfe de Californie Golfe de Californie Golfe de Californie Black sea Black sea Black sea Black sea Black sea Black sea Black sea Black sea Black sea Black sea Black sea Black sea Black sea Black sea Peru Peru Peru Peru Argentina Argentina Argentina Sea of Okhotsk Sea of Okhotsk Sea of Okhotsk Sea of Okhotsk Sea of Okhotsk Sea of Okhotsk Sea of Okhotsk Cariaco basin Cariaco basin Cariaco basin Cariaco basin Cariaco basin Cariaco basin Cariaco basin Cariaco basin Cariaco basin Cariaco basin Cariaco basin Cariaco basin Peru margin Peru margin Peru margin Peru margin Benguela margin Benguela margin Benguela margin Benguela margin Benguela margin Benguela margin Benguela margin Benguela margin Benguela margin Benguela margin Benguela margin Benguela margin Benguela margin Benguela margin Benguela margin Benguela margin Benguela margin Benguela margin Benguela margin Benguela margin Benguela margin Benguela margin Benguela margin Benguela margin Benguela margin Benguela margin Benguela margin Benguela margin Benguela margin Black and Azov sea Black and Azov sea Black and Azov sea Black and Azov sea Black and Azov sea Black and Azov sea Black and Azov sea Black and Azov sea Black and Azov sea Baltic sea Baltic sea Baltic sea Baltic sea Baltic sea Baltic sea Baltic sea Baltic sea Baltic sea Baltic sea Baltic sea Baltic sea Baltic sea Baltic sea Baltic sea Baltic sea Baltic sea Baltic sea Baltic sea Baltic sea Baltic sea Baltic sea Baltic sea Baltic sea Baltic sea Baltic sea Baltic sea Baltic sea Baltic sea Baltic sea Bornholm basin Bornholm basin Bornholm basin Bornholm basin Bornholm basin Bornholm basin Bornholm basin Bornholm basin Bornholm basin Bornholm basin Bornholm basin Bornholm basin Bornholm basin Bornholm basin Bornholm basin Bornholm basin Bornholm basin Bornholm basin Bornholm basin Bornholm basin Bornholm basin Bornholm basin Bornholm basin Bornholm basin Bornholm basin Bornholm basin Bornholm basin Bornholm basin Bornholm basin Bornholm basin Bornholm basin Bornholm basin Bornholm basin Bornholm basin Bornholm basin Bornholm basin Bornholm basin Bornholm basin Bornholm basin Bornholm basin Bornholm basin Bornholm basin Bornholm basin Bornholm basin Bornholm basin Bornholm basin Bornholm basin Bornholm basin Bornholm basin Bornholm basin Bornholm basin Bornholm basin Bornholm basin Bornholm basin Bornholm basin Bornholm basin Santa Barbara Basin Santa Barbara Basin Santa Barbara Basin Santa Barbara Basin Santa Barbara Basin Santa Barbara Basin Santa Barbara Basin Santa Barbara Basin Santa Barbara Basin Santa Barbara Basin Santa Barbara Basin Santa Barbara Basin Santa Barbara Basin Santa Barbara Basin Santa Barbara Basin Santa Barbara Basin Santa Barbara Basin Santa Barbara Basin Santa Barbara Basin Santa Barbara Basin Santa Barbara Basin Santa Barbara Basin Santa Barbara Basin Santa Barbara Basin Santa Barbara Basin Santa Barbara Basin Santa Barbara Basin Gulf of Aqaba Gulf of Aqaba Northen Red sea Northen Red sea Northen Red sea Northen Red sea Northen Red sea Northen Red sea Northen Red sea Northen Red sea Northen Red sea Northen Red sea Northen Red sea Northen Red sea Northen Red sea Northen Red sea Northen Red sea Northen Red sea Northen Red sea Northen Red sea Northen Red sea Andaman sea South China sea South China sea South China sea Bornholm Basin Bornholm Basin Bornholm Basin Bornholm Basin Bornholm Basin Bornholm Basin Bornholm Basin Bornholm Basin Bornholm Basin Bornholm Basin Aarhus Bay Aarhus Bay Aarhus Bay Aarhus Bay Aarhus Bay Aarhus Bay Aarhus Bay Aarhus Bay Aarhus Bay Aarhus Bay Aarhus Bay Aarhus Bay Aarhus Bay Aarhus Bay Aarhus Bay Baltic Sea Baltic Sea Baltic Sea Baltic Sea Baltic Sea Baltic Sea Baltic Sea Baltic Sea Baltic Sea Baltic Sea Baltic Sea Baltic Sea Baltic Sea Baltic Sea Baltic Sea Baltic Sea Baltic Sea Baltic Sea Baltic Sea Baltic Sea Baltic Sea Baltic Sea Baltic Sea Baltic Sea Baltic Sea Baltic Sea Baltic Sea Blake Bahama basin 4214 32.84 -10.78 1.10 Scharff (1980) Deep basin off Morocco Nielsen (2012) Baltic Sea 54.52 10.02 0.04 Hartmann and Nielsen (2012) Baltic Sea 54.52 10.02 0.07 Hartmann and Nielsen (2012) Baltic Sea 54.52 10.02 0.09 Hartmann and Nielsen (2012) Baltic Sea 54.52 10.02 0.12 Hartmann and Nielsen (2012) Baltic Sea 54.52 10.02 0.15 Hartmann and Nielsen (2012) Baltic Sea 54.52 10.02 0.18 Hartmann and Nielsen (2012) Baltic Sea 54.52 10.02 0.21 Hartmann and Nielsen (2012) Baltic Sea 54.52 10.02 0.24 Hartmann and Nielsen (2012) Baltic Sea 54.52 10.02 0.26 Hartmann and Nielsen (2012) Baltic Sea 9 45.47 37.33 0.05 Migdisov et al. (1974) Sea of Azov 9 45.47 37.33 0.24 Migdisov et al. (1974) Sea of Azov 9 45.47 37.33 0.81 Migdisov et al. (1974) Sea of Azov 46.15 36.63 0.02 Migdisov et al. (1974) Sea of Azov 46.15 36.63 0.06 Migdisov et al. (1974) Sea of Azov 46.15 36.63 0.22 Migdisov et al. (1974) Sea of Azov 46.15 36.63 0.92 Migdisov et al. (1974) Sea of Azov 45.90 35.25 0.03 Migdisov et al. (1974) Sea of Azov 100 56.33 20.37 0.05 Lein (1983) Baltic Sea 110 55.00 19.00 0.95 Lein (1983) Baltic Sea 476 34.36 -120.11 0.05 Kaplan et al. (1963) Santa Barbara Basin 1760 27.15 -0.23 Lein et al. (1976b) Gulf of California 110.87 2772 26.13 -110.41 0.45 Sweeney and Kaplan Gulf of California 476 34.36 -120.11 0.95 Kaplan et al. (1963) Santa Barbara Basin 1760 27.15 -0.23 Lein et al. (1976b) Gulf of California 110.87 (1980) 2772 26.13 -110.41 0.53 Sweeney and Kaplan Gulf of California 860 33.57 -118.47 0.05 Sweeney (1972) Santa Catalina Basin 1760 27.15 -0.55 Lein et al. (1976b) Gulf of California 110.87 (1980) 2772 26.13 -110.41 0.68 Sweeney and Kaplan Gulf of California 860 33.57 -118.47 0.21 Sweeney (1972) Santa Catalina Basin 1760 27.15 -0.55 Lein et al. (1976b) Gulf of California 110.87 (1980) 2772 26.13 -110.41 0.95 Sweeney and Kaplan Gulf of California 860 33.57 -118.47 0.43 Sweeney (1972) Santa Catalina Basin 1170 25.53 -0.63 Lein et al. (1976a) Gulf of California 110.55 (1980) 3361 24.70 -109.17 0.18 Sweeney and Kaplan Gulf of California 410 33.53 -118.86 0.05 Kaplan et al. (1963) Santa Catalina Basin 3260 26.35 -0.43 Lein et al. (1976a) Gulf of California 109.93 (1980) 3361 24.70 -109.17 0.25 Sweeney and Kaplan Gulf of California 859 33.75 -118.82 0.05 Kaplan et al. (1963) Santa Monica Basin 2860 23.78 -0.40 Lein et al. (1976a) Gulf of California 108.67 (1980) 3361 24.70 -109.17 0.35 Sweeney and Kaplan Gulf of California 859 33.75 -118.82 0.70 Kaplan et al. (1963) Santa Monica Basin 2860 23.78 -2.00 Lein et al. (1976a) Gulf of California 108.67 (1980) 3361 24.70 -109.17 0.55 Sweeney and Kaplan Gulf of California 1773 43.50 32.00 0.05 Vinogradov et al. (1962) Black Sea 1000 23.23 -0.90 Lein et al. (1976a) Gulf of California 107.12 (1980) 3361 24.70 -109.17 0.75 Sweeney and Kaplan (1980) Gulf of California 3361 24.70 -109.17 0.95 1179 41.58 29.33 0.05 Vinogradov et al. Black Sea 2760 0.50 -86.17 0.50 Migdisov et al. (1983) Costa Rica Pacific margin (1962) 2005 42.53 38.42 0.03 Migdisov et al. (1974) Black Sea 3120 1.98 -86.17 0.50 Migdisov et al. (1983) Barbados Sweeney and Kaplan Gulf of California 2005 42.53 38.42 0.61 Migdisov et al. (1974) Black Sea 3120 1.98 -86.17 1.20 Migdisov et al. (1983) Barbados (1980) 3361 24.70 -109.17 1.05 Sweeney and Kaplan (1980) Gulf of California 3361 24.70 -109.17 1.25 Sweeney and Kaplan Gulf of California 2075 43.87 36.90 0.03 Migdisov et al. (1974) Black Sea 2360 4.00 -86.17 0.50 Migdisov et al. (1983) Barbados 2008 43.83 33.67 0.05 Vinogradov et al. Black Sea 3330 2.53 -1.20 Migdisov et al. (1983) Atlanti Florida margin 101.08 (1962) (1980) 200 -44.63 167.55 0.03 Kaplan et al. (1960) Milford Sound 1704 42.00 40.18 0.05 Vinogradov et al. Black Sea The dataset gathered in table B.2 include a subsample of the compilation realized by Johnson Pyrite δ 34 S (‰) -39,29 -44,38 -40,96 -39,92 -33,82 -36,62 -33,47 -38,8 -45,11 -47,52 -33,48 -41,2 -42 -42,9 -43,4 -37,6 -39,7 -41,1 -43,8 -41,1 -41,5 -35,9 -41,5 Pyrite δ 34 S (‰) -42,8 -43,8 -43,2 -43,3 -44 -39 -41,9 -13,6 -13,5 -11,8 -16,4 -10,7 -17,3 -15,8 -38,4 -31,8 -33,5 -25,8 -21,2 -21 -21,3 -15,3 -16,8 -14,6 -14,1 -15 -17,6 Pyrite δ 34 S (‰) -14,8 -18 -10,5 -27,7 -1,6 -9,5 -11,4 -7,8 -27,1 5,4 -15,7 -24,6 -15,8 -9,3 -15,7 -4,4 10,1 -0,7 -19,4 -14,6 -13,8 -10,2 -8,7 -16,4 -17,1 -43,5 -42,6 Pyrite δ 34 S (‰) -32,2 -32 -32,7 -36,4 -31,8 -20,9 -38,2 -29,9 -36,5 -38,8 -37,1 -37,7 -1,4 -31,6 -34,3 -32,4 -18,2 -30,3 -9,4 -25,1 -36,1 -27,4 -34,7 -17,2 -36,3 -35,6 -28,8 Pyrite δ 34 S (‰) -34,5 -22,7 -17,3 -26,8 -16,3 -28,2 -40,2 -34,5 -39 -24,8 32,3 -32,3 -31,7 -28,3 -26,9 -22,5 -28,2 -24,6 -22 -27,6 -15,1 -17,6 -16,9 18,4 -34,3 -26,1 -14,1 Pyrite δ 34 S (‰) -8,2 -10,1 -13 -27,3 -29,2 -21,7 -7,8 -46,3 -6,5 -3,2 -6,6 -7,7 -42,9 -24,6 9,4 -11,5 -34,3 -34,5 -39,5 -15,1 -26,9 -7,9 -17,3 Pyrite δ 34 S (‰) -3,2 -7,9 -28,7 -42,9 17,6 -28,7 -38,4 -19,3 1,8 -26,3 -2,5 -13,2 -37,7 -27 -39,3 4,7 1,9 -37,5 -30,3 -22,2 -32 Pyrite δ 34 S (‰) -30,2 -3,3 -8 -29,4 -23,4 -24,2 -24,9 -32,1 -17 -31,4 -19,6 -19,8 -27,1 -24,6 -27,5 -12,7 -21,5 -20,6 -22,1 10,6 -25,5 Pyrite δ 34 S (‰) -21,1 -20,8 -35,8 0,1 -16,9 -36 -30,3 -10,3 3,3 -36,4 -21,9 -30,5 -28,7 -34 -36,1 -27,5 -21,1 -8,5 19,6 7,5 7,7 Pyrite δ 34 S (‰) -28,9 -20,9 -16,7 -31,3 -44,8 -36,5 -36 -36,3 -45,7 -46,6 -50,7 -50,8 -52 -51,5 -51,8 -52,4 -52,3 -52,4 -53,2 -51,8 -51,4 Pyrite δ 34 S (‰) -53,6 -52 -52,4 -52,2 -52,6 -51,4 -52 -47,1 -52,7 -46,6 -50,4 -50,9 -52,7 -52,7 -53,1 -51,2 -51,8 -53,2 -54,2 -52,9 -51,8 Pyrite δ 34 S (‰) -54,6 -52,9 -53,6 -53,5 -52,4 -52,9 -53,7 -52,7 -52,3 -52,6 -52,8 -52,4 -52,8 -53,4 -53,4 -52,3 -53,3 -52,5 -53,6 -53,8 -53,4 Pyrite δ 34 S (‰) -52,2 -52,3 -53,7 -52,5 -52,3 -53,3 -52,5 -53,7 -52,4 -53,1 -52,8 -52,4 -53 -51,4 -52,4 -53,2 -51,4 -53,2 -52,1 -51,9 -51,3 Pyrite δ 34 S (‰) -48,8 -48,8 -24,9 -47,5 -48,1 -46,9 -45,2 -43,8 -42,5 -38,9 -37,9 -38,9 -33,8 -30,1 -41,1 -47,7 -45,9 -46 -33,1 -45,2 -21,3 -10,7 -4,7 Pyrite δ 34 S (‰) -26,4 -19,6 -22,9 -13,6 -21,4 -12,4 -18,2 -0,8 4,5 -1,9 -9,2 -3,2 -5,2 0,8 -31,8 -19,6 -17,8 -7,2 -11,8 -6,3 -20,5 -29,1 -18,5 -28,1 -22,1 -20,8 -24,2 Pyrite δ 34 S (‰) -28,1 -32,2 -27,6 -24,1 -16,5 -13,9 -22,3 -27,8 -27,6 -19,8 -18,1 -25,2 -32,4 -26,8 -30,5 -29,2 -20,7 -18,1 -25 -28,1 -22,3 -18,6 -18,3 -29,1 -30,9 -22,7 -15,8 Pyrite δ 34 S (‰) -8,2 2,1 -23,9 -19,4 -23,8 -25,5 -27,6 -24,7 -14,3 -28,7 -27,8 -23,2 -25,6 -22,4 -18,5 -23,2 -20 -22,8 -15,4 -22,6 -16,7 -8,6 -3 -11,2 -6,7 -6,3 2,8 Pyrite δ 34 S (‰) -2,2 3 -7,3 33,1 10,2 24 26,6 1,6 0,3 11,2 -2,2 1,8 4,3 -2,5 -11,8 -13,7 -2,5 -7,5 -9,4 -16,3 -10,5 -18,9 -15,2 -20,9 -13,7 -16,4 -8,4 Pyrite δ 34 S (‰) -5,5 -16,2 -10 -12,5 19,9 -8,3 -14,5 -5,3 -4,4 0,9 -7,2 0,9 -2,8 -11,8 -3,4 -7,3 -6,7 -6,8 -2,7 Pyrite Pyrite δ 34 S (‰) Pyrite δ 34 S (‰) Pyrite δ 34 S (‰) Pyrite δ 34 S (‰) Pyrite δ 34 S (‰) Pyrite δ 34 S (‰) Pyrite δ 34 S (‰) Pyrite δ 34 S (‰) Pyrite δ 34 S (‰) Pyrite δ 34 S (‰) Pyrite δ 34 S (‰) Pyrite δ 34 S (‰) Pyrite δ 34 S (‰) Pyrite δ 34 S (‰) Pyrite δ 34 S (‰) δ 34 S (‰) 2.37 -26.7 -31 -35 -23.8 -28 -27 -39 6.97 6.78 6.21 4.93 7.24 -36.51 -36.2 -34.41 -28.85 -19.53 -37.44 -33.92 -29.68 -30.9 -26.9 -27.3 -31 -29.9 -49.1 -48.9 -49.7 -18.2 -10.6 -17.2 -23.7 -19.3 -23 -21.1 -32.6 -33.2 -31.9 -33.8 -33.3 -32.7 -32.9 -32.7 -32.3 -32.9 -31.3 -31.7 -29 -27.9 -35.1 -33.3 -29 -32.1 -22 -22 -22.4 -32.1 -29.5 -29.9 -35.7 -23.3 -36 -41.5 -42.5 -42.6 -2.2 -5 -4.2 -4.2 -2.3 -4.9 -1.8 -4.3 -24.5 -12.8 -23.5 -13.5 -34.5 -33.2 -34.2 -27 -8.7 -3 -9 -1.5 -1.6 -12.2 -18.1 -10.2 -12.2 -28.5 -33.3 -32.2 -33.5 -31.8 -31.9 -31.1 -30.2 -29 -28.5 -28.4 -14.9 -23.2 -25.7 -23.8 -21.8 -22.2 -23.5 -23 -23 -23.5 -22.2 -21.9 -21.8 -23.3 -22.9 -21.9 -9.5 -24.6 -28.3 -26.4 -22.9 -19.7 -25.1 -23.3 -26.5 -21.9 -20.1 -10.4 -16.1 -26.9 -31.6 -31.2 -31.1 -28.0 -29.8 -29.0 -26.0 -28.1 -26.6 -33.4 -32.4 -27.7 -35.8 -28.5 -35.8 -36.0 -32.5 -26.0 -34.8 -34.8 -33.6 -31.6 -32.0 -32.5 -39.6 -37.3 -40.5 -38.0 -39.8 -40.4 -32.1 -39.3 -32.7 -40.3 -39.4 -39.8 -39.5 -37.9 -36.3 -37.4 -31.1 -36.4 -39.6 -37.1 -24.5 -25.7 -20.9 -25.8 -28.3 -32.1 -32.2 -33.1 -35.3 -35.6 -36.7 -33.8 -35.5 -35.7 -37.7 -36.6 -37.9 -38.2 -34.8 -33.8 -31.3 -27.8 -26.9 -31.2 -29.3 -33.8 -32.6 -39.02 -41.76 -36.89 -29.87 -27.19 -27.56 -26.24 -28.18 -27.23 -25.08 -26.10 -26.79 -26.83 -24.22 -26.17 -26.94 -24.40 -27.17 -26.79 -28.24 -27.36 -29.6 -6.0 -12.3 -16.4 -35.5 -36.0 -32.4 -26.1 -28.8 -31.0 -36.1 -39.1 -36.9 -37.3 -31.8 -31.9 -30.9 -30.3 -27.4 -28.3 -25.7 -28.0 -29.0 -29.5 -27.6 -26.7 -25.4 -28.9 -24.6 -20.2 -22.1 -23.4 -23.4 -22 -22.4 -20.1 -19.3 -23.8 -27.6 -29 -27.3 -25.9 -26.1 -27.2 -26.7 -26.6 -27.1 -27 -25.1 -24.9 -26.3 -25.8 -25 -20 -26 -26 -43.6 -15.1 -26.1 -19.4 -21.3 -23.2 -24.2 -24.2 -24 -22.8 -22.1 -22.9 -23.5 -20.3 -23.6 -4.2 6.2 13.7 10.8 11 -20 -26.6 -26.7 -19.6 -17.3 -26.5 -18.2 -28 -31 -21.5 -30.2 -30.7 -17.3 -35.1 -35 -14.2 -29.9 -23.8 -23.6 -10.7 -28 -23.2 -15.1 -24 -23.2 -33.7 -39 -22.7 -24 -48.7 -26.3 -44.7 -29.4 -22 -44.7 -20.2 -17 -16.1 -48.9 -29.2 -33.2 -44 -25.46 -33.7 (1962) 4 45.90 35.25 0.10 Migdisov et al. (1974) Sea of Azov 6.5 54.75 14.00 0.09 Lein (1983) Baltic Sea -18 54.75 14.00 0.51 Lein (1983) Baltic Sea -15.2 54.75 14.00 1.01 Lein (1983) Baltic Sea -14 45.03 36.57 0.05 Migdisov et al. (1974) Black Sea -25.9 130 57.93 21.08 0.04 Lein (1983) Baltic Sea -24.2 130 57.93 21.08 0.20 Lein (1983) Baltic Sea -25 130 57.93 21.08 0.50 Lein (1983) Baltic Sea -26 200 -44.63 167.55 0.03 Kaplan et al. (1960) Milford Sound -28.9 140 23.38 2163 42.50 39.00 0.03 Vinogradov et al. Black Sea and Adkins 2022, the rest constitute a compilation of data from the PANGEA database and -30.6 (1962) -106.95 0.70 Lein et al. (1976a) Gulf of California -27 1030 32.74 2003 42.50 34.30 0.03 Vinogradov et al. Black Sea from a recent literature review of the available datased for surficial sediment iron sulfide -27.4 (1962) -117.55 0.05 Kaplan et al. (1963) San Diego Trough 3.1 1030 32.74 117.55 0.95 Kaplan et al. (1963) San Diego Trough -28.8 226 28.63 -112.33 0.00 (1980) San Pedro Martir Basin -20.6 226 28.63 -112.33 0.09 (1980) San Pedro Martir Basin 1529 -50.65 -46.10 0.70 Scharff (1980) Argentina basin -18.4 -21.9 Sweeney and Kaplan 2003 42.50 34.30 0.41 (1962) Black Sea -25.3 Vinogradov et al. Sweeney and Kaplan 2003 42.50 34.30 0.18 (1962) Black Sea -24.9 Vinogradov et al. -2003 42.50 34.30 0.03 Vinogradov et al. (1962) Black Sea -27.4 isotopic composition.

  Aloisi et al. / Geochimica et Cosmochimica Acta xxx (2022) xxx-xxx 7

Please cite this article in press as: Aloisi G., et al. The geochemical riddle of ''low-salinity gypsum" deposits. Geochim. Cosmochim. Acta (2022), https://doi.org/10.1016/j.gca.2022.03.033 from values approaching those of modern Mediterranean Sea salt works

Table 1

 1 Summary of crystal-bound water isotopic composition and fluid inclusion salinity of Primary Lower Gypsum and Upper Gypsum deposits.

	Basin	Section	Gypsum parent water				Fluid inclusion	Reference
									salinity
			d 18 O H2O (‰ vs SMOW)	dD H2O (‰ vs SMOW)
			Min	Max Av. (n b )	Min	Max Av.	Min Max Av.(n b )
	Caltanissetta	EM	3.64 5.78 4.71 (31)	16.99 28.05	22.11 -	-	-	This work
		SE	1.03 3.60 2.43 (14)	À1.51 13.88	5.36 -	-	-	This work
		GI	À2.62 2.51 0.48 (6)	-22.44 13.78	À0.77 -	-	-	This work
	Vena del Gesso	Mo/MT	0.99 4.34 2.74 (35)	0.91 22.07	11.49 24.5 78.2 52.2 (6) This work
	Piedmont	BA	0.28 3.15 1.76 (45)	1.61 23.48	11.86 2 a	77 a	16 a (7)	Natalicchio et al.
												(2014)
												and this work
		AR	À3.47 2.38 À0.47	À31.96 17.59	À7.39 -	-	-	This work
					(27)
		PO	À7.00 2.48 À1.47	À56.43 10.64 À14.57 -	-	-	This work
					(72)
	Sorbas	RA	À2.06 3.21 0.83 (25)	À17.04 15.03	À0.36 18.8 50.7 32.3	Evans et al. (2015)
											(11)
	Catanzaro	MQ	-	-	-	-	-	-	1.8	65 f	27 (24)	Costanzo et al. (2019)
	Trough
	Modern salina	Cabo de Agua	6.3	7.6	7.42 (5)	29.6	38.62	33.13 -	-	-	Evans et al. (2015)
		Cagliari	6.2	7.3	-	-	-	-	-	-	-	Longinelli (1979)
		Aigues-Mortes	7.1 c 8.7 c 7.90 (8)	27.0	37.6	32.07 -	-	-	Fontes (1966)Fontes
												and Gonfiantini (1967)
		Salin-de	4.5 d 6.4 d 5.13 (3)	18.9 d 29.5 d	22.50 -	-	-	Pierre (1982a)
		Giraud	3.7 e 4.6 e 4.03 (3)	12.4 e 17.4 e	14.30 -	-	-	Pierre (2018)

Notes: EM-Eraclea Minoa; SE-Santa Elisabetta; GI-Gibliscemi; MT-Monte Tondo; BA-Banengo; AR-Arnulfi; PO-Pollenzo; RA-Rio de Aguas; MQ-Marcellinara Quarry.

a Fluid inclusion salinity was measured by

[START_REF] Natalicchio | Did Late Miocene (Messinian) gypsum precipitate from evaporated marine brines? Insights from the Piedmont Basin (Italy)[END_REF]

on gypsum samples from the Banengo and Moncucco sections. b Indicates the number of macroscopic gypsum samples investigated. c Values reported by Pierre (1982a,b). d mm-thick, seasonal gypsum crust sampled in July 1976. e Pluri-annual, cm-thick gypsum crusts subject to seasonal cycles of dissolution/re-cristallisation by meteoric waters, sampled in March 1978. f Disregards one outlier salinity vaule at 115 g kg À1 . 8 G. Aloisi et al. / Geochimica et Cosmochimica Acta xxx (2022) xxx-xxx Please cite this article in press as: Aloisi G., et al. The geochemical riddle of ''low-salinity gypsum" deposits. Geochim. Cosmochim. Acta (2022), https://doi.org/10.1016/j.gca.2022.03.033

Table 2 )

 2 : (i) riverine input from Mediterranean Sea (principally Nile and Rhone) and Black and Caspian Seas 10 G. Aloisi et al. / Geochimica et Cosmochimica Acta xxx (2022) xxx-xxx

Please cite this article in press as: Aloisi G., et al. The geochemical riddle of ''low-salinity gypsum" deposits. Geochim. Cosmochim. Acta (2022), https://doi.org/10.1016/j.gca.2022.

03.033

Table 2

 2 Chemical and isotopic composition of the Mediterranean, Black and Caspian seas and of the major rivers that feed them.

	12		G. Aloisi et al. / Geochimica et Cosmochimica Acta xxx (2022) xxx-xxx
		Water	[Sr 2+ ]	87 Sr/ 86 Sr	[Ca 2+ ]	[SO 4 2À ]	d 34 S SO4	References
		discharge	(lM)		(mM)	(mM)	vs CDT
		(km 3 /yr)
	Mediterranean Sea rivers
	Nile	83	2.68	0.706-	0.775	0.542	-	Brass (1976)Gerstenberger et al.
				0.7071				(1997)
	Rhone	54	5.94	0.7087	1.770	0.479	-	Gaillardet et al. (1999)
	Black Sea rivers
	Danube	207	2.76	0.7089	1.473	0.660	7.6	Gaillardet et al. (1999)Burke
								et al. (2018)
	Dniester	11	-	-	1.725	0.319	5.5	Burke et al. (2018)
	Dnieper	54	2.5	0.7084	1.105	0.356	5.5	Gaillardet et al. (1999)Palmer
								and Edmond (1989)Burke et al.
								(2018)
	Don	28.9	2.5	0.7084	2.073	0.953	5.8	Gaillardet et al. (1999)Palmer
								and Edmond (1989)Burke et al.
								(2018)
	Caspian Sea rivers
	Volga	239.7	5.48	0.708083	1.1	0.67	7.4	Clauer et al. (2000)Kroonenberg
								et al. (2005)Mekhtieva and
								Rabinovich (1975)
	Ural	8.8	-	-	-	-	5.2	Mekhtieva and Rabinovich
								(1975)
	Western rivers	33.6	18.8	0.707795	-	-	-	Clauer et al. (2000)
	Southern rivers	10.2	9.7	0.708293	-	-	-	Clauer et al. (2000)
	Water masses
	Atlantic Ocean	-	87	0.709006	10	28.9	22	McArthur et al. (2001)Berner
								and Berner (2012)Masterson
								et al. (2016)
	Caspian Sea	-	114	0.708183	7.95	30	8.7	Clauer et al. (2000)Mekhtieva
								and Rabinovich (1975)
	Notes: Western Caspian rivers include Sulak, Samur, Kura and Terek.

Please cite this article in press as: Aloisi G., et al. The geochemical riddle of ''low-salinity gypsum" deposits. Geochim. Cosmochim. Acta (2022), https://doi.org/10.1016/j.gca.2022.

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  . G. Aloisi et al. / Geochimica et Cosmochimica Acta xxx (2022) xxx-xxx 13

Please cite this article in press as: Aloisi G., et al. The geochemical riddle of ''low-salinity gypsum" deposits. Geochim. Cosmochim. Acta (2022), https://doi.org/10.1016/j.gca.2022.03.033 luminescence (supplementary Fig.

Table 3

 3 Mass balance estimate of the percentage of re-crystallized gypsum needed to explain gypsum parent water d 18 O values. Average d 18 O of the crystallization water of gypsum precipitated by evaporation of Mediterranean waters under summer (July) evaporative conditions (see Fig. 2f for corresponding water isotope composition). G. Aloisi et al. / Geochimica et Cosmochimica Acta xxx (2022) xxx-xxx 17 Paratethys formed a surface layer that underwent little or no mixing with the underlying Mediterranean highsalinity water. The brackish surface layer filled Mediterranean marginal basins where it formed low-salinity gypsum by evaporation. The key to this scenario is that modern Caspian Sea water, taken to represent Paratethys water during the Messinian Salinity Crisis

	Section	Messinian gypsum a	Meteoric H 2 O	Pure evaporitic gypusm e	Recristallized fraction
		Min d 18 O	Max d 18 O	d 18 O	d 18 O	Max%	Min%
	Eraclea Minoa	7.14	9.28	À7.00 b	6.5	21.2	5.3
	S. Elisabetta	4.53	7.10	À7.00 b	6.5	40.5	21.5
	Gibliscemi	0.88	6.01	À7.00 b	6.5	67.6	29.6
	Monte Tondo	0.99	4.34	À8.00 c	6.5	62.1	39.0
	Banengo	3.78	6.65	À9.00 d	6.5	40.1	21.6
	Arnulfi	0.03	5.88	À9.00 d	6.5	64.3	26.6
	Pollenzo	À3.50	5.98	À9.00 d	6.5	87.1	25.9
	Sorbas	1.44	6.71	À5.11 d	6.5	73.7	28.3
	Note:
	a Oxygen isotope composition of gypsum crystallization water.

b Liotta et al. (2013). c Giustini et al. (2016). d Evans et al. (2015). e Please cite this article in press as: Aloisi G., et al. The geochemical riddle of ''low-salinity gypsum" deposits. Geochim. Cosmochim. Acta (2022), https://doi.org/10.1016/j.gca.2022.03.033

  suggest, a two-way water exchange flow with the Mediterranean Sea existed, then the Paratethys would have rapidly acquired normal marine Ca 2+ /Cl À and SO 4 2À /Cl À ratios by the onset of the Messinian Salinity Crisis. These considerations exclude that the Paratethys source of freshwater carried a dissolved 18 G. Aloisi et al. / Geochimica et Cosmochimica Acta xxx (2022) xxx-xxx

Please cite this article in press as: Aloisi G., et al. The geochemical riddle of ''low-salinity gypsum" deposits. Geochim. Cosmochim. Acta (2022), https://doi.org/10.1016/j.gca.2022.

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Table 4

 4 Yearly cycle of hydrological and biogeochemical processes that lead to the formation of one mm-thick gypsum lamina in marginal basins.

	Precession	Season	Reservoir Process
	Minimum precession = humidity Humid and arid Surface	Stratification decreases the transport of nutrients from deep water layer to surface layer, NO 3 -based PP stops
				N 2 -based primary production until P depletion, then oligotrophic surface layer (process 1):

  Please cite this article in press as: Aloisi G., et al. The geochemical riddle of ''low-salinity gypsum" deposits. Geochim. Cosmochim. Acta (2022), https://doi.org/10.1016/j.gca.2022.03.033 hydrological conditions, intense continental runoff kept the water column permanently stratified throughout the annual cycle. Stratification limited the transport of nutrients from the deep waters to surface waters; NO 3

			G. Aloisi et al. / Geochimica et Cosmochimica Acta xxx (2022) xxx-xxx
			CaSO 4
	Humid	All	Same stratified conditions as during minimum precession
			21

  ), it is not surprising that, 22 G. Aloisi et al. / Geochimica et Cosmochimica Acta xxx (2022) xxx-xxx in the face of a quasi-constant d 34 S SO4 , the d 18 O values of dissolved sulfate show a range of about 9‰ vs SMOW (Fig. 3). In support of this interpretation, the gypsum deposits from the Eraclea Miona section, that have the highest d 18 O SO4 values, are also those that have highest d 18 O H2O (Fig. 2b). In summary, the d 34 S SO4 and d 18 O SO4 compositions of Messinian Salinity Crisis low-salinity gypsum deposits are consistent with an efficient H 2 S reoxidation in waters of variable d 18 O H2O within the CSB mat.

Please cite this article in press as: Aloisi G., et al. The geochemical riddle of ''low-salinity gypsum" deposits. Geochim. Cosmochim. Acta (2022), https://doi.org/10.1016/j.gca.2022.

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  ; (2) Ceratolithus acutus is very rare also in fully marine open-ocean sediments (e.g. Di Stefano and Sturiale, 2010); (3) despite being rare in the late Messinian Mediterranean, this species has never been documented together with other long-range taxa, generally predominant in the assemblage, in Stage 3 deposits (see discussion in

Table 1

 1 Compilation of the 87 Sr/ Sr ratios from the Eraclea Minoa section plotted in Fig.2. Asterisks indicate that the analyzed samples are not assigned by the original authors to one of the sub-units distinguished in this work.

	Cycle	Sample	Sample type	Sub-unit	Stratigraphic height	87 Sr/ 86 Sr	Error	Reference
					(m)
	7	EM 8-11	Calcite (C. Agrigentina)	D	233,5	0.70864	6E-06	Grossi et al. (2015)
	7E M 8 -9	C a l c i t e ( C. Agrigentina)	D	232	0.70863	8E-06	Grossi et al. (2015)
	7E M 8 -7	C a l c i t e ( C. Agrigentina)	D	231	0.708647	1E-05	Grossi et al. (2015)
	7E M 8 -3	C a l c i t e ( C. Agrigentina)	D	228,5	0.708704	7E-06	Grossi et al. (2015)
	7	EM-44	Gypsum (Selenite)	B*	227*	0.708732	1E-05	García-Veigas et al. (2018)
	7	G7-8	Gypsum (Selenite)	C	226.2	0.708710	8.40E-06	This work (Bristol)
	7	G7-6	Gypsum (Selenite)	C	225.7	0.708704	1.08E-05	This work (Bristol)
	7	G7-4	Gypsum (Selenite)	C	225	0.708730	1.40E-05	This work (Bristol)
	7	G7-1	Gypsum (Selenite)	C	224.2	0.708720	6.50E-06	This work (Bristol)
	7	EM-41	Gypsum (Balatino)	B*	223*	0.708729	1E-05	García-Veigas et al. (2018)
	6	EM 7-10	Calcite (C. Agrigentina)	D	206	0.708637	1.1E-05	Grossi et al. (2015)
	6E M 7 -9	C a l c i t e ( C. Agrigentina)	D	205	0.708602	2E-06	Grossi et al. (2015)
	6E M 7 -5	C a l c i t e ( C. Agrigentina)	D	202.8	0.708629	7E-06	Grossi et al. (2015)
	6E M 7 -2	C a l c i t e ( C. Agrigentina)	D	201	0.708552	9E-06	Grossi et al. (2015)
	6	EM 6"-20	Calcite (C. Agrigentina)	D	197	0.708666	1E-05	Grossi et al. (2015)
	6	EM 6"-17	Calcite (C. Agrigentina)	D	191	0.708510	7E-06	Grossi et al. (2015)
	6	EM 6"-14a	Calcite (C. Agrigentina)	D	188	0.708511	5E-06	Grossi et al. (2015)
	6	EM 6"-8	Calcite (C. Agrigentina)	D	182.3	0.708685	1.1E-05	Grossi et al. (2015)
	6	EM 6"-7	Calcite (C. Agrigentina)	D	172.3	0.708609	7E-06	Grossi et al. (2015)
	6	EM 6"-5	Calcite (C. Agrigentina)	D	170.3	0.708654	9E-06	Grossi et al. (2015)
	6	EM 6"-3	Calcite (C. Agrigentina)	D	168.3	0.708608	9E-06	Grossi et al. (2015)
	6E M 6 ' -2	C a l c i t e ( C. Agrigentina)	D	163.3	0.708663	7E-06	Grossi et al. (2015)
	6E M 6 ' -2	C a l c i t e ( C. Agrigentina)	D	162.8	0.708729	1E-05	Grossi et al. (2015)
	6	EM 6'-22b	Calcite (C. Agrigentina)	D	161.3	0.708619	5E-06	Grossi et al. (2015)
	6E M 6 ' -1	C a l c i t e ( C. Agrigentina)	D	160.3	0.708671	5E-06	Grossi et al. (2015)
	6E M 6 ' -1	C a l c i t e ( C. Agrigentina)	D	159.3	0.708602	7E-06	Grossi et al. (2015)
	6E M 6 ' -1	C a l c i t e ( C. Agrigentina)	D	158.3	0.708658	9E-06	Grossi et al. (2015)
	6E M 6 ' -9	C a l c i t e ( C. Agrigentina)	D	157.3	0.708614	5E-06	Grossi et al. (2015)
	6E M 6 ' -6	C a l c i t e ( C. Agrigentina)	D	156.3	0.708685	5E-06	Grossi et al. (2015)
	6E M 6 ' -3	C a l c i t e ( C. Agrigentina)	D	155.3	0.708675	1E-05	Grossi et al. (2015)
	6E M 6 ' -1	C a l c i t e ( C. Agrigentina)	D	154.8	0.708612	1.1E-05	Grossi et al. (2015)
	6	EM-39	Gypsum (Balatino)	C*	154.2*	0.708724	1E-05	García-Veigas et al. (2018)
	6	EM-84	Gypsum (Selenite)	C	154.2	0.708735	1.2E-05	This work (Glasgow)
	6	EM-78	Gypsum (Selenite)	C	152.5	0.708744	1.4E-05	This work (Glasgow)
	6	EM-76	Gypsum (Selenite)	C	151	0.708740	1.3E-05	This work (Glasgow)
	6	EM-69	Gypsum (Selenite)	C	149	0.708723	1.2E-05	This work (Glasgow)
	6	EM-64	Gypsum (Balatino)	B	148.5	0.708747	1.3E-05	This work (Glasgow)
	6	EM-63	Gypsum (Balatino)	B	148	0.708738	1.3E-05	This work (Glasgow)
	6	EM-33	Gypsum (Balatino)	B*	148*	0.708719	1E-05	García-Veigas et al. (2018)
	5	EM-31	Gypsum (Selenite)	C*	133.8*	0.708751	1E-05	García-Veigas et al. (2018)
	5	EM-58	Gypsum (Selenite)	C	133.8	0.708748	1.3E-05	This work (Glasgow)
	5	EM-56	Gypsum (Selenite)	C	132.8	0.708748	1.2E-05	This work (Glasgow)
	5	EM-54	Gypsum (Selenite)	C	132.2	0.708756	1.2E-05	This work (Glasgow)
	5	EM-52	Gypsum (Selenite)	C	131.3	0.708776	1.4E-05	This work (Glasgow)
	5	EM-38	Gypsum (Selenite)	C	130.6	0.708756	1.3E-05	This work (Glasgow)
	5	EM-36	Gypsum (Selenite)	C	129.8	0.708765	1.6E-05	This work (Glasgow)
	5	EM-34	Gypsum (Selenite)	C	129	0.708770	1.4E-05	This work (Glasgow)
	5	EM-33	Gypsum (Selenite)	C	128.2	0.708764	1.4E-05	This work (Glasgow)
	5	EM-31	Gypsum (Selenite)	C	127.5	0.708777	1.4E-05	This work (Glasgow)
	5	EM-25	Gypsum (Balatino)	B	126.6	0.708789	1.4E-05	This work (Glasgow)
	5	EM-22	Gypsum (Balatino)	B*	126.6*	0.708747	1E-05	García-Veigas et al. (2018)
	4	EM-20	Gypsum (Selenite)	C*	115.8*	0.708769	1E-05	García-Veigas et al. (2018)
	4	G4-10	Gypsum (Selenite)	C	115.8	0.70876	6.21E-06	This work (Bristol)
	4	G4-8	Gypsum (Selenite)	C	114.5	0.708765	5.23E-06	This work (Bristol)
	4	G4-6	Gypsum (Selenite)	C	113	0.708767	9.79E-06	This work (Bristol)
	4	G4-4	Gypsum (Selenite)	C	111.5	0.708767	8.32E-06	This work (Bristol)
	4	G4-1	Gypsum (Selenite)	C	109.5	0.708786	3.89E-06	This work (Bristol)
	4	EM-13	Gypsum (Balatino)	B*	108*	0.708783	1E-05	García-Veigas et al. (2018)
	3	EM4-19	Calcite (C. Agrigentina)	D	105	0.708725	1.2E-05	This work (Bristol)
	3	EM4-17	Calcite (C. Agrigentina)	D	104	0.7088	3E-05	This work (Bristol)
	3E M 4 -7C a l c i t e ( C. Agrigentina)	D	97	0.708729	1E-05	Grossi et al. (2015)
	3	EM4-0K	Calcite (C. Agrigentina)	D	94	0.708827	6.30E-06	This work (Bristol)
	3	EM4-00a	Calcite (C. Agrigentina)	D	88	0.70873	1.2E-05	This work (Bristol)
	3	EM-12	Gypsum (Selenite)	C*	88*	0.708773	1E-05	García-Veigas et al. (2018)
	3	G3-11	Gypsum (Selenite)	C	87.70	0.708769	6.069E-06	This work (Bristol)
	3	G3-9	Gypsum (Selenite)	C	87.05	0.708778	6.364E-06	This work (Bristol)
	3	G3-7	Gypsum (Selenite)	C	86.4	0.708796	7.27E-06	This work (Bristol)
	3	G3-4	Gypsum (Selenite)	C	85.5	0.708779	5.60E-06	This work (Bristol)
	3	G3-3	Gypsum (Selenite)	C	85.15	0.708780	6.18E-06	This work (Bristol)
	3	G3-1	Gypsum (Selenite)	C	84.5	0.708785	6.12E-06	This work (Bristol)
	3	EM-7	Gypsum (Balatino)	B*	84*	0.708777	1E-05	García-Veigas et al. (2018)
	2	EM-6	Gypsum (Selenite)	C*	48*	0.708780	1E-05	García-Veigas et al. (2018)
	2	G2-11	Gypsum (Selenite)	C	47.75	0.708780	5.83E-06	This work (Bristol)
	2	G2-9	Gypsum (Selenite)	C	47.2	0.708776	7.05E-06	This work (Bristol)
	2	G2-7	Gypsum (Selenite)	C	46.65	0.708774	8.25E-06	This work (Bristol)
	2	G2-5	Gypsum (Selenite)	C	46.1	0.708781	4.77E-06	This work (Bristol)
	2	G2-3	Gypsum (Selenite)	C	45.55	0.708780	7.67E-06	This work (Bristol)
	2	G2-1	Gypsum (Selenite)	C	45	0.708778	6.47E-06	This work (Bristol)
	2	EM-2(a)	Gypsum (Balatino)	C*	45*	0.708775	1E-05	García-Veigas et al. (2018)
	1	Cum-5	Gypsum (Balatino)	B	2.7	0.708800	5.5E-06	This work (Bristol)
	1	Cum-3	Gypsum (Balatino)	B	1.5	0.708810	5.6E-06	This work (Bristol)

G.Aloisi et al. / Geochimica et Cosmochimica Acta xxx (2022) xxx-xxx

(caption on next page)F.Andreetto et al.

or elemental sulfur: S • ). The arrows represent the sulfur fluxes described in Supplementary section 5. (For interpretation of the colors in the figure, the reader is referred to the web version of this article.)

ACKNOWLEDGEMENTS

We thank Felice Tirabasso and Matthias Reiman (Knauf) for granting access and permission to sample the Castellina Marittima quarry (Tuscany, Italy). We thank Stefano Marabini and Saint Gobain for granting access and permission to sample the Monte Tondo quarry (Vena del Gesso, Italy). We thank Christophe Thomazo for carrying out the d 18 O SO4 analyses at the Universite ´de Bourgogne (Dijon). This work was supported by the French TELLUS-Marges program and by the European Union's Horizon 2020 research and innovation programme under the Marie Skłodowska-Curie grant agreement No 765256 (ETN SaltGiant). This work benefitted from very useful comments provided by the associated editor Adrian Immenhauser and two anonymous reviewers.

Acknowledgments

We thank the entire SALTGIANT community for many profitable workshops that inspired the development of this manuscript. This research was supported by the project SALTGIANT-Understanding the Mediterranean Salt Giant, a European project which has received funding from the European Union's Horizon 2020 research and innovation program, under the Marie Sklodowska-Curie [grant agreement No 765256]. We greatly thank the two reviewers Domenico Cosentino and William Ryan and the editor Alessandra Negri for the fruitful comments provided that led to a substantial improvement of the manuscript.

Acknowledgements

A special thanks goes to Jamie Lewis and Carolyn Taylor from the Bristol Isotope Group Facilities, who ran the Sr isotope analyses during the Covid pandemic, therefore giving continuity to the work. We thank Elsa Gliozzi and Domenico Cosentino for providing unwashed marl samples and the washed residues of Grossi et al. (2015), which helped to expand the ostracods' Sr dataset, Tilly Bouten for assistance with SEM analyses and Frits Hilgen, Paul Meijer, Rob Govers, Mhina de Vos for the numerous and fruitful discussions which this manuscript greatly benefited from. Antonio Caruso and Isabel Sanchez-Almazo are thanked for joining the field campaign in the early noughties. We thank the Editor and two anonymous reviewers for their thorough comments and suggestions, that greatly improved the original version of the manuscript. This research was supported by the project SALTGIANT-Understanding the Mediterranean Salt Giant, a European project which has received funding from the European Union's Horizon 2020 research and innovation program, under the Marie Sklodowska-Curie [grant agreement No 765256]. The Royal Society provided funding for the analytical and fieldwork that took place in the early noughties through a Wolfson Dorothy Hodgkin Fellowship to RF.

Acknowledgements

The authors would like to thank Christophe Thomazo for δ 18 O SO4 analyses as well as Ivan Jovovic and James Bradley for their precious suggestions and assistance in this work. Aaron Meilijson and the anonymous reviewer are thanked for their constructive review and comments that greatly improved this manuscript. We thank Stefano Marabini and Saint Gobain for granting access and permission to sample the Monte Tondo quarry (Vena del Gesso, Italy). We are grateful to all SALTGIANT ESRs, PIs and POs for the numerous discussions and comments during workshops, courses, and fieldtrips. This research is carried out under the SALT-GIANT ETN, a European project funded by the European Union's Horizon 2020 research and innovation program under the Marie Sklodowska-Curie grant agreement number 765256.

Context of the PhD -The SaltGiant ETN The present thesis has been undertaken at the Institut de Physique du Globe de Paris (IPGP) under the supervision of Giovanni Aloisi. My PhD work was also carried out within the SaltGiant European Training Network (ETN), a project funded by the European Union's Horizon 2020 Research industry, geo-economics of the Mediterranean region and the history of oceanography. viii The SaltGiant ETN (2018-2022) funded 15 PhD projects -of which my PhD is one -across various disciplines including geology, geophysics, geochemistry, microbiology and the social sciences. To address its objectives, the 15 PhD projects are organized into 4 work packages

Data availability

The compilation of strontium, sulfate and hydrogen isotope data plotted in figure 9, as well as some of the available, and here not (graphically) shown, oxygen isotope values is accessible in separate excel spreadsheets (Supplementary material 1).

Data availability

All data used for the research are present in the article.

Declaration of Competing Interest

The authors declare that they have no known competing financial interests or personal relationships that could have appeared to influence the work reported in this paper.

APPENDIX A. SUPPLEMENTARY MATERIAL

Supplementary data to this article can be found online at https://doi.org/10.1016/j.gca.2022.03.033.

Declaration of Competing Interest

The authors whose names are listed immediately below certify that they have NO affiliations with or involvement in any organization or entity with any financial interest (such as honoraria; educational grants; participation in speakers' bureaus; membership, employment, consultancies, stock ownership, or other equity interest; and expert testimony or patent-licensing arrangements), or non-financial interest (such as personal or professional relationships, affiliations, knowledge or beliefs) in the subject matter or materials discussed in this manuscript.

Appendix A. Supplementary data

Supplementary data to this article can be found online at https://doi. org/10.1016/j.earscirev.2021.103577.

F. Andreetto, R. Flecker, G. Aloisi et al. Earth and Planetary Science Letters 595 (2022) 1177672 and/or 3 on ocean biogeochemistry occurred in the Pliocene once Mediterranean outflow was re-established.

CRediT authorship contribution statement

Federico Andreetto conducted the field work in 2020 and 2021, prepared the samples for 87Sr/86Sr analyzes and paleontological observations and wrote the manuscript. Wout Krijgsman, Rachel Flecker and Vanni Aloisi supervised the research and contributed to the writing of the manuscript. Alan Maria Mancini identified the calcareous nannofossils and contributed to the discussion. Laetitia Guibourdenche contributed to the discussion. Stephanie de Villiers analyzed the 87Sr/86Sr ratios from gypsums V and VI in Cambridge (2002). All authors revised the work critically and approved the manuscript.

Declaration of competing interest

The authors declare that they have no known competing financial interests or personal relationships that could have appeared to influence the work reported in this paper.

Contents lists available at ScienceDirect

Earth and Planetary Science Letters Net fractionation factors 34 ε net and 33 λ net calculated for the gypsum (massive, banded, and branching selenite) and marl samples using the isotope compositions of the sulfate ion in equilibrium with gypsum and the sulfur isotope composition of iron sulfides (dots and crosses) compared with ranges of 34 ε and 33 λ values obtained in pure cultures of sulfur metabolizing microorganisms (shaded areas). For marls samples, where no gypsum precipitates occur, an isotope composition of the sulfate ion equal to the Messinian ocean value (Masterson et al., 2016) was set, based on the assumption that sulfides from marls were produced through the reduction of oceanic sulfate. See Supplementary section 2. Supplementary section 9 provides full details on the modeling and results of the various biogeochemical scenarios.

In all scenarios (i) to (v), the entire range (from 0 to 1) of f ox was explored ( f ox refers to the fraction of H 2 S produced by MSR being reoxidized; see Supplementary section 5 and 6 for a full description of the model equations and parameters). The isotope fractionation factors are fixed for all processes except for MSR, for which the full ε MSR and λ MSR ranges have been explored (see section 4.3.2 for further explanation).

Cryptic sulfur cycling during the formation of VdG gypsum layers

The best fit of the model results to the analytical dataset (Fig. 8)

is obtained with f msr = 0.8 (representing the fraction of sulfur entering the system as SO 4 and subsequently reduced by MSR, see Supplementary section 7), and scenario (iv), i.e. the partial oxidation of sulfide to elemental sulfur by chemolithotrophs and subsequent disproportionation of elemental sulfur (Figs. 7C,8). See Supplementary section 8 for further discussion on the reoxidative processes involved in sulfur cycling. In this best fit model run, 90% of the isotopic data in both diagrams (δ 34 S SO4 -33 S SO4 , Fig. 8A; 34 ε net -33 λ net , Fig. 8B) plot within the area defined by 0.85 < f ox < 1. The combined sulfate and sulfide isotope dataset can therefore only be reproduced by considering that: (1) at least 80% of the net incoming sulfate is reduced microbially to H 2 S rather than directly precipitated as gypsum ( f msr > 0.8) (Fig. 8), and (2) nearly all of the MSR-produced H 2 S is reoxidized (0.85 < f ox < 1) (Fig. 8).

We suggest that the relatively high ε MSR and f ox values inferred for gypsum compared to lower values for marls indicates the occurrence of alternations between drastically different biogeochemical environments associated with periods of marl and gypsum deposition (Fig. 7,8). These biogeochemical changes were controlled by orbitally driven climatic oscillations triggering changes in hydrological regime. During marl deposition (i.e. precession minima), freshwater inputs were higher, leading to water column stratification, less re-oxidation of hydrogen sulfide and an efficient preservation of organic matter (Natalicchio et al., 2019;Sabino et al., 2021) (Fig. 1a, 7a). Conversely, during gypsum precipitation (precession maxima), intense reoxidation of hydrogen sulfide prevailed (Fig. 7b), requiring at least seasonal mixing of the water column.

The large majority of sulfur (>80%) forming gypsum was first reduced from sulfate to sulfide by MSR, possibly stored as elemental sulfur and then almost completely reoxidized back to sulfate (>85%) resulting in low net production of H 2 S (Fig. 7B,C) and confirming the existence of a "cryptic" sulfur cycle during gypsum formation.

Implications for carbon cycling and gypsum formation

To our knowledge, this is the first time that gypsum deposits are proven to be formed from sulfate sourced almost entirely from 10

Declaration of competing interest

The authors declare that they have no known competing financial interests or personal relationships that could have appeared to influence the work reported in this paper.

Appendix A. Supplementary material

Supplementary material related to this article can be found online at https://doi .org /10 .1016 /j .epsl .2022 .117676.