Dr Véronique Gardien

location Direction of deposition

Keywords: Brillouin Spectroscopy, Evaporites, Gypsum, Dead Sea, Sound Velocity, Palaeoclimate reconstructions . . . . . . . . . . . . . . . . . . . . . . . . . Brillouin spectroscopy, Dead Sea Basin, Ćuid inclusions, marine isotope stage (MIS) 5e, palaeotemperature, seasonal halite facies

Evaporites such as salt and gypsum are common minerals in, for example, oceans, sabkhas, caves, desertic regions and lakes. Evaporite sequences can be useful climate archives: physical and chemical conditions of formation of the host mineral can be retrieved from water droplets (Ćuid inclusions; FIs) trapped in the crystal during its growth. Recently, a novel method has been developed to measure sound velocity in FIs and to determine their entrapment temperatures in host halite crystals. It is based on Brillouin Spectroscopy (BS), which measures small frequency shifts in inelastically scattered light. Can BS be used to generate high-resolution temporal reconstructions of climate for regions and time intervals with evaporite deposition? To address this question, we applied BS on halite sections from the Dead Sea Basin (DSB). The DSB is a unique study site, because halite layers were deposited annually (varves) during drier and warmer interglacial periods (Marine Isotope Stages; MIS 7c and MIS 5e).

BS measurements on halite FIs reveal maximum temperature differences of 4 °C between bottom and uppermost intervals of individual layers during both interglacial intervals. Palaeotemperatures from the oldest halite section (ca 215 ka; MIS 7c) further highlight a decadal warming of the lake bottom waters during summer (+3 °C) and winter (+4 °C) periods along the core section. These results also show annual winter lake temperature Ćuctuations throughout all layers, whereas summer lake temperatures remained constant for six years, followed by annual variations until the end of halite deposition. A possible connection between varve thicknesses and summer water temperatures is also discussed. These results have major implications for future studies on layered halite sequences at larger geographical and temporal scales.

The question remains to which extent BS can be applied to other minerals, such as the evaporitic sulphate mineral gypsum. Experiments were carried out using BS analysis on capillary solutions with various NaCl concentrations as well as on synthetic gypsum. This method was also applied to natural gypsum samples from (high) saline environments (e.g., Tunisian sabkha, Aral Sea) and from outcrop dating back to the Messinian Salinity Crisis in the Mediterranean realm. i Gypsum entrapment temperatures scarcely reproduce the expected values for synthetical samples of known water temperatures. Experiments carried out with natural gypsum samples, on the other hand, provide a larger quantity of T data from FIs, suggesting that natural gypsum samples might serve for temperature studies using BS. Further work is needed to conĄrm this.

Developing the methodology on gypsum is further motivated by the sensitivity of sound velocity on salinity, as observed in capillary solution experiments and in biphasic FIs in natural gypsum. Thus, sound velocity can be used as a high throughput salinity proxy in evaporites making Brillouin spectroscopy a promising tool for future palaeosalinity reconstructions.

RESUMÉ

Le sel ou le gypse sont des minéraux courants des séries évaporitiques se formant dans des contextes variés comme les océans, les sebkhas, les grottes, les régions désertiques et les lacs. Lors de leur croissance ces minéraux piègent des microgouttelettes dŠun Ćuide (inclusions Ćuides : IFs) ayant préservé les conditions physiques et chimiques à partir desquels les cristaux ont poussé. De fait, ses séquences dŠévaporites sont des archives climatiques potentielles. Récemment, une nouvelle méthode a été développée pour mesurer la vitesse du son dans les inclusions Ćuides dŠhalite aĄn de déterminer leur température de piégeage. Cette méthode, la spectroscopie Brillouin (BS) mesure les petits décalages de fréquence (le shift Brillouin) de la lumière diffusée inélastiquement.

La question majeure que je me suis posé au cours de cette thèse est de savoir si la BS peut être utilisée pour reconstruire à haute résolution lŠévolution de la température dans des régions et pour des intervalles de temps où se forme un dépôt dŠévaporite ? Pour répondre à cette question, nous avons choisi dŠétudier un site unique, le bassin de la mer Morte où des dépôts dŠhalite se sont déposés en couches annuelles (varves) pendant les périodes interglaciaires (stades isotopiques marins ; MIS 7 et MIS 5e). Un forage ICDP effectué en 2011 au centre de la Mer Morte a permis dŠavoir accès à cette séquence sédimentaire qui couvre les 200 derniers ka. Pour les deux intervalles interglaciaires, les mesures de BS sur les FI dŠhalite révèlent une différence de température maximales de 4 °C entre la base et le sommet de chaque varve. Les paléotempératures de la section dŠhalite la plus ancienne (environ 200 ka ; MIS 7) soulignent en outre un réchauffement décennal des eaux du fond du lac pendant les périodes estivales (+3 °C) et hivernales (+4 °C). De plus, pour tous les dépôts dŠhalite nous observons des Ćuctuations annuelles de la température du lac en hiver. Pour les dépôts estivales les températures sont restées constantes pendant 6 ans suivies de variations annuelles jusquŠà la Ąn du dépôt dŠhalite. Nous montrons quŠil existe un lien possible entre lŠépaisseur des varves et les températures estivales de lŠeau. Ces résultats majeurs ont des implications sur les futures études des séquences halitiques à des échelles de temps et dŠespace plus grandes.

iii La question subsidiaire à laquelle jŠai tenté de répondre au cours de ma thèse est de savoir si la BS peut être appliquée à dŠautres minéraux présents dans les séquences évaporitiques, comme le gypse. Des expériences ont été menées en utilisant lŠanalyse BS sur des solutions capillaires avec différentes concentrations de NaCl ainsi que sur du gypse synthétique. Cette méthode a également été appliquée à des échantillons de gypse naturel provenant dŠenvironnements hyper salins tels quŠune sebkha tunisienne, la mer dŠAral et des dépôts Messiniens dans le domaine méditerranéen. Si les températures de piégeage du gypse synthétique obtenus avec la BS reproduisent à peine les valeurs attendues en revanche les données sur des échantillons de naturels sont encourageant et suggèrent que le gypse naturel est un bon candidat pour les reconstructions de paleotempératures à lŠaide de la BS.

Le développement méthodologique sur le gypse à permit de mettre en évidence la sensibilité de la vitesse du son par rapport à la salinité de la solution, faisant de la spectroscopie Brillouin un outil prometteur pour les futures reconstructions de paléosalinité.

Mots clés: Spectroscopie Brillouin, Évaporites, Gypse, Mer morte, Vitesse du son

Reconstructions paléoclimatiques ACKNOWLEDGEMENT I am grateful to be part of the Make Our Planet Great Again program by Campus France, and for the Ąnancial support for more than three years. Furthermore, I wish to express my sincere thanks to the University of Lyon 1, the Institute of Light and Matter, the Doctoral School, and to our team ŠLiquides et InterfacesŠ for all pandemic-related extensions obtained (especially Emmanuel Martelat for his approval to extend the contract once again, and all other team members).

I would like to thank all the jury members for their willingness to review and evaluate this work.

I express my sincere thanks to my two supervisors Véronique Gardien and Frédéric Caupin for their support in difficult times, for their patience, for their time spent on administrative and experimental issues, as well for all the nice moments we shared together apart the institute. Not a month went by without a problem to solve, but we handled it pretty well collectively as a group. I am proud to have worked with you! I have appreciated the great help in private and professional matters by Daniel Ariztegui and Philippe Sorrel, who provided me with moral support and encouragement whenever required, for providing samples from Lake Petén Itzá and the Aral Sea, and for their scientiĄc advice to publish my Ąrst article. This project could not have been realized without Hedi Oberhänsli, Peter Zavialov, Francesco Dela Pierre and Marcello Natalicchio, and Markus Schwab, who provided valuable sampling material to be studied during the last three years. Thank you very much to all of you! Furthermore, I thank Giovanni Aloisi and all members of SaltGiant, as well as the colleagues from Palermo for organising and inviting me for the amazing Ąeld trip to Sicily in 2019.

Furthermore, I thank Thierry Alboussière for his help, time, and patience with tax issues, and Jean Colombani for providing the gypsum powder, without which much v of this work could not have been realised. Special thanks to Emanuel Guillerm for the intensive training phase on the instruments, his constant assistance and help to deal with scientiĄc problems, and for introducing me to the Dead Sea subject.

Thanks to Olivier for the blind trust in the beginning and the exciting time in Craponne (greetings to Julien, Maxime and Misca). Thanks to Marie, with whom I spent most of my time with a wide variety of (quirky) personalities, and who showed me many beautiful places in France.

The time in Lyon was also marked by short but intense encounters (greetings to Amanda, Cecilia and Massimo, Marion, and all former Ćatmates), as well as beautiful moments among fellow countrymen (Theo, Johannes and especially Anne, our visits to the pub together were always great, and thanks for showing me the good taste of IPA beers). I thank Chenglong Ge for exciting Mölki games and for showing me chicken paws and delicious Chinese ŠwineŠ. A shout out goes to Wiem, my fellow sufferer in the lab, our lunch breaks and Wizard games were always a great distraction to me. I couldnŠt have completed this work without my great friends from the office who provided many bright spots away from university. Thanks to Team Mont Blanc for numerous video game nights, mineral exhibitions, hikes, football sessions, etc (Antoine and Raphaël), and to Laëtitia who became my closest friend here. You were there for me all the time, in funny and bad moments, which was immensely important to me. I will miss you, our funny conversations in the office, and our activities around Lyon. (Lyell and Deshayes, 1830). Hutton was one of the major drivers of modern geology in the 18th century, because he hypothesised that sediments at the Earth surface must be much older than previously thought.

Further, he developed the pioneering concept of "uniformitarianism" that assumes quite the same natural laws and principles prevailing today and in the past, thus, being responsible for the formation of geological units (Hutton, 1788;Hutton, 1795). After HuttonŠs death in 1797, it was Charles Lyell who published three volumes of "Principles of Geology" (1830-1833) in which he deĄned for the Ąrst time the Tertiary period, and more importantly, his ideas about the origin of glaciers and moraines (Lyell, 1837). Around at the same time, similar observations were made by Louis Agassiz with the outcome of his glaciation theory (Agassiz, 1840;Flint, 1968), which convinced also the famous geologist William Buckland who had already described glacial traces in Scotland many years before (Buckland, 1824).

It was Svante Arrhenius, a physical chemist from Sweden, who mathematically calculated the effect of atmospheric carbon dioxide (CO 2 ) concentrations on global temperatures in 1896. Based on his calculations, he pointed out that the global climate should rather move towards decreasing global temperatures to end up in a new ice age (Arrhenius, 1896). Afterwards, the presence of periodic orbital variations was described mathematically in terms of the rotation of the earth axis ("precession"; 23,000 years), the degree of the EarthŠs spin axis ("obliquity"; 41,000 years), and the EarthŠs distance to the sun ("eccentricity"; 120,000 years) by Milankovitch, 1930. This had led to the hypothesis that the EarthŠs climate varied over time in the past, controlled by the amount of incoming solar radiation.

The discovery of stable isotopes (Thomson, 1912) with the development of the Ąrst mass spectrometer by Aston (1920) opened a new world in particular for geoscientists later.

Another pioneering work was done by Giaque and Johnston who measured different masses of oxygen atoms apart from the known mass of 16 which was the discovery of oxygen isotopes 17 O and 18 O ( Giauque and Johnston, 1929a;Giauque and Johnston, 1929b). Oxygen isotopic analysis (δ 18 O) of minerals is still by far the most widely used proxy in palaeoenvironmental studies. Probably as important as oxygen was the discovery of the heavy isotope hydrogen ŠdeuteriumŠ with 2 H (Urey and Teal, 1935). After WWII, the techniques developed for previously military purposes were widely released for civilian and scientiĄc use. For example, Harold Urey, being involved in the Manhattan Project to participate in building the nuclear bomb, focused during the post-war period on the extension of oxygen isotopes to geological studies. He reported the mass-dependent fractionation processes of isotopes and particularly between carbonates and water (Urey, 1947), further he provided the Ąrst evidence for a temperature-dependency of δ 18 O based on Cretaceous calcareous fossils (Urey, 1948;Urey et al., 1951). At the same time, the Ąrst linear palaeotemperature scale for inorganic carbonates has been proposed by McCrea (1950), including both the temperature T in °C and the oxygen composition ∆ 18 O in :

T = 18 -5.4 * ∆ 18 O (1.1)
Two years later, a palaeothermometer for marine organic carbonates was published based on both, laboratory-grown (lab-grown) organically precipitated carbonate species and collected marine calcareous shells, with the isotopic composition reported in δ 18 O (Epstein et al., 1951):

T = 11.88 -5.91 * δ 18 O (1.2)
Two years later, this equation had to be revised because organic matter from shells has inĆuenced the previous isotopic analyses (Epstein et al., 1953):

T = 16.5 -4.3 * δ 18 O + 0.14 * δ 18 O (1.3)
Furthermore, there were indications of a weak linear relationship between the salinity (ranging from 33 to 38 ) and the abundance of the heavier oxygen isotope 18 O in water samples (Epstein et al., 1951).

These fundamental works have led Cesare Emiliani, an Italian micropalaeontologist working in UreyŠs geochemistry labŠ in Chicago, to analyse benthic foraminifera (single-celled marine organisms with calcareous shells). First, material obtained from paciĄc bore cores was used to determine the oxygen isotopic composition to reconstruct ocean bottom water temperatures in four samples from the Oligocene (Emiliani and Edwards, 1953), continued by a long-term Pleistocene reconstruction covering the last 290 ka with the deĄnition of isotope stages and the observation of Ćuctuating marine temperatures using 1.3 (Emiliani, 1955;Emiliani, 1961;Emiliani, 1970). Emiliani should became one of the pioneer of modern palaeooceanography, especially in view of his results, showing cyclical variations during the Pleistocene similar to the mathematical predictions by Milankovitch in 1930. It is worth noting that, once again, the outcome of scientiĄc cooperation among interdisciplinary Ąelds (Šcrosscutting conceptŠ) enabled the geological importance of EmilianiŠs work (among others), by the discovery of the radiocarbon dating method ( 14 C) (Libby et al., 1949). However, 14 C dating is limited in geological timescales to ca 50 ka due to its short half-life of 14 C (5730 ±40 years; Godwin, 1962).

Therefore, and due to the immense technological advances, various ŠabsoluteŠ dating methods have been developed (e.g., U-Th, U-U, U-Pb, Rb-Sr, K-Ar, etc.) besides ŠrelativeŠ dating methods (biostratigraphy, Marine Isotope Stages, tephrochronology, palaeomagmatism, etc.). For instance, palaeoclimatic events during the last Ąve million years could be assigned to magnetic reversals based on palaeomagnetism which had also broad implications for the community and the understanding of EarthŠs climate (Hays and Opdyke, 1967;A. Cox et al., 1964). Shortly afterwards, Shackleton, 1969 revisited the Ąndings of Emiliani, 1961 for the last interglacial warming period (ca 130 -80 ka) and introduced isotopic sub-stages for Marine Isotope Stage 5 (MIS;a,b,c,d,e). However, the exact cause of EarthŠs temperature variations during MIS 5 remained ambiguous because the observed temperature rises/drops in the records were asymmetric (Shackleton, 1969).

Finally, a study dealing with marine cores in the Southern Hemisphere impressively proved the connection between orbital variations, as predicted by Milankovitch, and the occurrence of global ice ages in the geological record (Hays et al., 1976).

Another major outcome of this study was a model predicting a long-term Northern Hemispheric glaciation over the next several thousand years without anthropogenic effects (Hays et al., 1976).

After the Ąrst discovery of related increased atmospheric CO 2 concentrations and rising global temperatures in the late 1930Šs (Callendar, 1938), the 1980Šs then were marked by a rapid global temperature rise, and further studies pointed out the dramatic impact of atmospheric CO 2 concentrations on global temperatures and rising sea-levels due to polar ice melting (e.g., Hansen et al., 1981;Gornitz et al., 1982). A study reported variations in atmospheric CO 2 concentrations from the Vostok ice core, Antarctica, for the past 160 ka (Barnola et al., 1987). Although it showed that such variations are linked to interglacial-glacial climate states, it also revealed that atmospheric CO 2 concentrations have never been such high than in the 1980Šs (338-350 ppm from 1980-1989). In 1989, the Intergovernmental Panel on Climate Change (IPCC) was founded to get a scientiĄc-based global state of the EarthŠ climate by combining numerous scientiĄc disciplines. Since then Chapter 1. Introduction six assessment reports have been published in 1990, 1995, 2001, 2007, 2014, and 2022, all of them generally have increased the warnings towards drastic changes not only in climate and extinction of species, but also on inhabitable regions for humans. Modern climate change is also termed Šglobal warmingŠ because it takes into account the mean global atmospheric temperature rise relative to the pre-industrial period (1850), but in fact this temperature rise does not affect all regions equally (IPCC, 2022). The latest report by IPCC (2022) has once again revelead that future mean temperatures will rise more at higher latitudes at the poles than at lower latitudes, but that regions at lower latitudes are not equally affected due to non-linear responses of complex oceanic/atmospheric circulation patterns. Moreover, reconstructed CO 2 over the past three million years suggested that Šby 2024/2025 levels of atmospheric CO2 will be higher than any point in the last 3.3 million yearsŠ (De La Vega et al., 2020).

Therefore, palaeoclimate studies by using geological records and proxy methods are still essential to provide detailed data sets from remote regions where monitoring data are missing. Although an extensive database is already available, there is still a substantial need for additional proxy methods to better investigate speciĄc climate archives. This in turn supports climate models to predict future changes of important climatic parameters such as temperature or salinity.

The main limitation of terrestrial palaeoclimate research is, however, the lack of accessible, suitable, continuous, and/or datable climate archives and proxies.

Nevertheless, the most powerful method is to combine data sets obtained from multiple proxy methods for models (Lelieveld et al., 2012). Evaporitic minerals, for example, are commonly present in arid environmental regions where other climate archives are absent, which makes them a potential gap Ąller for speciĄc dry periods of the Earth history.

Chapter 1. Introduction

Climate archives: Evaporites

Commonly, chemical salty minerals that form (sub)aqueously in arid or semiarid climate settings under solar-driven evaporation processes from a water body are deĄned as "evaporites" and make up a group of 60 major minerals (Warren, 2016). Evaporites form in marine, continental (including lakes, caves), and coastal environments. The (sea)water chemistry (i.e., ion composition) controls which evaporitic mineral forms with respect to the prevailing supersaturation state of each mineral (Warren, 2016). The composition of seawater is fairly homogeneous worldwide, with small anomalies (e.g., in the eastern Mediterranean), affecting the order of mineral precipitation due to solubility in marine environments:

• 1. Carbonates • 2. Gypsum • 3. Halite • 4. Potash minerals
The chemical composition of water bodies in continental regions depends mainly on regional geology, water balance, or even microbacterial activity, which makes it mostly very different from seawater. This also changes the order of the Ąrst forming minerals in nonmarine environments:

• 1. Carbonates • 2. Gypsum • 3. NaSO 4 , NaCO 3 , MgSO 4 • 4. MgCl 2 , CaCl 2 • 5. NaCl
Both lists represent the most common evaporitic sequences from the sediment record and modern observations (Warren, 2016), but it might differ in hyperarid regions with unique water chemistries (e.g., Dead Sea).

Main groups of evaporites are constituted by the major cations on EarthŠs surface, e.g., Ca 2+ , Mg 2+ , K + , Na + , Fe (Holser, 1979). Therefore, evaporites are quite diverse by incorporating either two (e.g., NaCl, "halite"), three (e.g., CaSO 4 , "anhydrite"), four (e.g., Mg 3 B 7 O 13 • Cl, "boracite") or even Ąve elements (e.g.,

Chapter 1. Introduction NaCaB 5 O 9 • 5 H 2 O
, "proberite"). However, despite the diversity, the dominating minerals are carbonates, sulphates, and salts. In most cases, their timing of formation depends on the degree of supersaturation with respect to a particular salty mineral (Warren, 2016). The supersaturation depends on the solubility of an individual mineral and thus the nucleation starts for each mineral differently according to the amount of evaporated water molecules of a water body. For example, the solubility of NaCl in 1 L of water is 360 g at 25 °C [START_REF] Haynes | Crc handbook of chemistry and physics[END_REF], so that the Ąrst NaCl crystal precipitates at a concentration >360 g/L. With an increase in (elemental) concentration of any solution, brine, or water body, the heat capacity decreases whereas the density increases (Warren, 2016).

Other processes than solar-driven evaporation to precipitate evaporites can involve biological processes or temperatures close to negative. For example, cryogenic salts like mirabilite (Nelson and Thompson, 1954), hydrohalite (Drebushchak et al., 2019), antarcticite (Kimura and Kai, 1984), and epsomite form at temperatures close to or below the freezing point of the liquid phase (Warren, 2016).

Both the temporal and spatial local distribution of evaporites are neither limited to space nor time in Earth history. The oldest evaporitic sequences have been reported from the Onega Basia, Russia, and dated to the Paleoproterozoic Era, that is, 2 billion years ago (Blättler et al., 2018). Further studies conĄrmed the occurrence of evaporites during the Proterozoic [START_REF] Spear | Analyses of Ćuid inclusions in Neoproterozoic marine halite provide oldest measurement of seawater chemistry[END_REF]Lindsay, 1987;[START_REF] Kah | Geochemistry of a 1.2 Ga carbonate-evaporite succession, northern Baffin and Bylot Islands: implications for Mesoproterozoic marine evolution[END_REF]. Besides the marine realm being a major accommodation space for carbonates, sulphates, and salts, they also form in caves, in saline pans, coastal environments, sabkhas, deserts, and in lakes.

Therefore, non-altered evaporites are potentially excellent climate archives that can be used for speciĄc geological periods.

Halite

Being part of the chlorides group, the mineral ŠhaliteŠ (or Šrock saltŠ) is composed of sodium chloride (NaCl) and is assigned to the isometric crystal system, thus, NaCl crystals are cube-shaped (a = 5.64 Å, α = 90°) (Figure and Macedo, 1996).

NaCl is globally distributed over time and space, and its suitability for palaeoclimate studies has been already reported for several locations (e.g., Lowenstein et al., 1999;Kiro et al., 2016;Schubel and Lowenstein, 1997;Hardie and Lowenstein, 2004).

Another beneĄt for experimental work is its high solubility (360 g/L), which allows to produce a vast amount of material in a day or two with a sufficiently amount of trapped Ćuid inclusions (Figure 1.1 B) (Rigaudier et al., 2012). Further it was

shown that the composition of stable isotopes of both oxygen and hydrogen of extracted FIs can be used as a proxy for isotopic trajectories of water evaporation, and to estimate temperature and wind velocity (Rigaudier et al., 2011). As pointed out later in this work in more detail, this early work on salty inclusions in NaCl has paved the way for extended palaeoclimatic applications of salt. Data obtained from material from the Dead Sea basin provided evidence for accurate temperature reconstructions using Brillouin Spectroscopy for progressive studies (Guillerm, 2019;Guillerm et al., 2020;Brall et al., 2022).

A B

Gypsum

The mineral gypsum is a member of the calcium sulfate group (CaSO 4 ) which holds in its crystal lattice up to 20% H 2 O as molecular water (Šcrystal waterŠ). The crystal system is monoclinic (a = 5.67 Å, b = 15.15 Å, c = 6.28 Å, and β = 113°50Š; Aronson et al., 1983). Gypsum is the most abundant sulfate mineral on EarthŠs surface and an important industrial product. The stability of gypsum is limited to less than 42 °C, a threshold above which it starts slowly to dehydrate to the hemihydrate form (ŠbassaniteŠ). However, as time is relative, the actual transformation takes place at higher temperatures of >70-90 °C. After complete dehydration it transforms further into anhydrite with a change into the orthorhombic crystal system (Figure 1.2). However, the inverse is possible by re-hydrating anhydrite, and to re-transform back into the stable gypsum phase. Hence, natural geological gypsum records that have been exposed on EarthŠs surface could have suffered from several transformation processes after deposition.

The abiotic growth mechanisms of gypsum minerals have long been undetermined due to the complexity of incorporated ions into the molecular crystal structure.

Based on in-situ observations in nucleation experiments of calcium sulfate crystals (e.g., He et al., 2018;Van Driessche et al., 2012), a more complex and non-universal nucleation pathway has been stated (Van Driessche et al., 2019). (Dolomieu, 1792). The Ąrst detailed description of such inclusions ("Ćuid inclusions"; FIs), has been published for a variety of minerals with the important suggestion that probably all crystals contain liquid cavities (Brewster, 1823). Further, it was then Sorby, 1858 who expanded the scientiĄc knowledge about Ćuid inclusions, e.g. by categorising quantitatively their sizes that can be still seen as valid today: calculated from his sizes in inches, his scale ranged from <2.5 µm (very small), ≤25.4 µm (moderately small), ≤254 µm (moderately large), to more than 254 µm (very large). During this thesis, the relative Ćuid inclusion size categories will be deĄned as very small (<2 µm), small (2-5 µm), intermediate

(5-15 µm), and large (>15 µm).

Trapped Ćuid inclusions are often mimicking the crystal system of its host mineral by negative shapes. For example, FIs in a cubic NaCl crystal are themselves cubicshaped (Figure 1.3 A, although other forms can also occur, e.g. as rectangularelongated forms (Figure 1.3 B). FIs that have been trapped during crystal growth in a speciĄc layer at similar depth, aligned along crystal growth zones, are termed ŠFluid inclusions assemblagesŠ (FIA) (Roedder, 1984). Due to the large complexity of numerous varieties of FIs in minerals at the EarthŠs surface, in the crust or mantle, this thesis will be only focussing on NaCl (ŠhaliteŠ) and CaSO 4 • 2 H 2 O (ŠgypsumŠ) minerals. For detailed information about analysis, studies, occurrences, or pitfalls, the reader is referred to Roedder (1984), Samson et al. (2003), [START_REF] Fall | How Precisely Can the Temperature of a Fluid Event be Constrained Using Fluid Inclusions?[END_REF].

Pristine or primary FIs (non-altered) are trapped as monophasic liquids at the Earth subsurface at ambient pressure (e.g., minerals forming in arid sabkhas, salt pans, or in the upper water layer in lakes) or at higher pressure (e.g., in metamorphic zones, or deeper water layer in lakes). Owing to major improvements in analytical techniques in the 20th century, studies of minerals containing FIs have been widely extended since the 1960Šs (e.g., Roedder, 1962;Roedder, 1963). The main beneĄt to study FIs is their phase behaviour with changing pressure and temperature, which allows to determine the relative elemental compositions (Roedder, 1972).

Moreover, the usage of microthermometry to determine maximum homogenisation temperatures (T h ), at which the vapor bubble of a two phase FI (liquid+vapor) vanishes, has been considered as a reliable tool to identify crystal formation temperatures in halite (Lowenstein et al., 1999. Fall andBodnar (2018) Further studies applied FIs as a palaeoenvironmental archive to diverse geological records, e.g., the occurrence of living halobacteria in FIs (Norton and Grant, 1988), crystal formation temperatures (Lowenstein et al., 1999;Guillerm, 2019), hydrothermal ore deposits (Wilkinson, 2001), the composition of FIs (Burke, 2001), DNA extraction from FIs that showed long-living cycles of halophilic archaea and bacteria in Dead Sea sediments (Thomas and Ariztegui, 2019), or δ 18 O variations in speleothem FIs (Griffiths et al., 2010).

Chapter 1. Introduction

Temperature proxy in this work

Despite decades of research in the Ąeld of geosciences (including palaeoclimate research), fundamental misconceptions about climate archives and climate proxies can still be found in many media and publications. A climate archive serves as material with physical, chemical, or biological properties which can be both quantitatively and qualitatively analyzed and measured by instruments, and thus providing information about past environmental conditions of a sample specimen (Jones and Mann, 2004). A climate proxy is both measurable and calibrated against an instrumental standard. Climate archives are solids (minerals, rocks, or layers) and are comparable to a computer hardware: tree rings, speleothems, ice cores, ocean sediments, lake sediments, and in more detail, (a)biogenic carbonates, quartz veins, or evaporites. On the other hand, climate proxies are like the "software" of a solid system by describing (and explaining) the nature of the solid, and by deĄning its environmental formation conditions. Hence, Ćuid inclusions are here deĄned as the ŠinternalŠ climate archive, whereas the host mineral would be the ŠexternalŠ archive. The derived property of FIs, here the temperature, is thus the climate proxy.

Analytical capabilities in the natural sciences have expanded rapidly since the technological advances of the 20th century, allowing a range of global climate archives to be studied. Hence, marine climate archives have been increasingly used to reconstruct paleoclimate. Although such oceanic bore cores are costly, they are available virtually everywhere, and they also cover the climate history of several million years. However, since continental archives are also important for understanding complex atmospheric processes and are both far less expensive and easier to sample, various temperature proxies, among others, have been developed. A frequent problem in continental archives such as caves, tree rings, lakes, Ćuvial sediments, cores, or outcrops, are among others erosion, limited temporal resolution, or the lack of directly measurable temperature proxies.

Huge advances during the past two decades have been reported for carbonates in temperature reconstructions using clumped isotope thermometry. This method, an extension of the classical isotope mass spectrometry, beneĄts from the "clumping" effect of relatively heavier isotopes 13 C and 18 O when CaCO 3 precipitates in (sub)aquatic solutions (Eiler, 2007;Eiler, 2011). The formation of heavier isotope bounding is directly related to temperature, and most important, not dependent on the oxygen isotope composition of the parent water. Hence, this technique circumvents the long-term issue of unknown palaeo-seawater oxygen compositions when applying the classical stable isotope method (δ 13 C; δ 18 ). The applicability of clumped isotopes has been shown on various carbonate species from nearly all Chapter 1. Introduction environments, e.g., caves (Affek et al., 2008), travertines and tufa (Kele et al., 2015), or soils (Kelson et al., 2020).

However, even such a seemingly impeccable technique has its limitations, because continental carbonates do not form everywhere and at all times, for example in lakes only under certain conditions with low sedimentation rates. Evaporites, on the other hand, Ąll this gap, at least in saline lakes, so that distinct periods could be reconstructed paleoclimatically for speciĄc regions such as the Dead Sea or arid sabkhas.

A common tool to reconstruct palaeotemperatures in evaporites has long been the microthermometry method that is based on the observation of the timing when a vapor bubble disappears (ŠhomogenizesŠ). It can be used to determine the salinity of samples, and thus is a well-established technique to obtain important information about the parent brine of the respective host mineral. It has been used in many geological domains, e.g., for pyrite (Kouzmanov et al., 2002), sulĄdes [START_REF] Zaw | Microthermometry and chemical composition of Ćuid inclusions from the Mt Chalmers volcanic-hosted massive sulĄde deposits, central Queensland, Australia: implications for ore genesis[END_REF], petroleum inclusions (Grimmer et al., 2003), pegmatites [START_REF] Beurlen | Fluid inclusion microthermometry in Be-Ta-(Li-Sn)-bearing pegmatites from the Borborema Province, northeast Brazil[END_REF] and salts (Lowenstein et al., 1999).

However, there are few limitations towards low-temperature minerals that commonly form at the EarthŠs surface at low-pressure. Recent experiments in synthetically-produced halite FIs have shown a signiĄcant offset (ca 5 °C) to the known crystal growth temperatures (Guillerm, 2019;Lowenstein et al., 1999).

Contrary to that, results obtained from a previously developed palaeothermometer (Brillouin spectroscopy, BS) gave evidence to overcome the relative large uncertainty by studies on pure water FIs in quartz (Mekki-Azouzi et al., 2015).

The BS method has further been adapted in detail for natural NaCl samples, and offered insights into lake-level variations in the Dead Sea basin (Guillerm, 2019).

However, there are potentially further studies dealing with sedimentary records from the Dead Sea possible, since the halite successions form an analogue to modern-day seasonal halite formation in the Dead Sea (Ben Dor et al., 2019;Sirota et al., 2017). Moreover, BS has never been tested on gypsum FIs in particular, which leaves a question unanswered if such abundant subsurface minerals might serve as an accurate climate archive when in combination with BS.

The major important advance of BS is the direct measurement of crossing temperatures derived from both monophasic and biphasic Ćuid inclusions along a controlled temperature-pressure path in the laboratory. It was shown that a small amount (ca 20 FIs per sample) suffices to measure accurately halite samples within a short period of time (Guillerm et al., 2020). Therefore, this approach has been extended to another evaporitic mineral, gypsum, and to seasonal halite layers from the Dead Sea basin. Since gypsum is not only a mineral but also a Chapter 1. Introduction common marine/ continental climate archive, the BS technique might help to reveal palaeoenvironmental conditions such as temperature for geological records.

History of salinity deĄnitions

The term "salinity" aims to describe the "saltiness" of a water body or solution, and thus refers generally to the amount of dissolved ions in water. It has undergone various deĄnitions throughout the past since its Ąrst deĄnition (Forchhammer, 1865). Around two decades later, [START_REF] Dittmar | Report on researches into the composition of ocean-water collected by HMS Challenger during the years 1873-76[END_REF] published Ąrst his Ąndings of mainly constant seawater salt concentrations on a global scale. From 1900 on, "chlorinity" has been used as Primary Standard of seawater describe the salinity of a solution which has been deĄned "as the weight in grams of chlorides obtained in 1 kg of sea water" in parts per thousand (Cl℃), being proportional to the mass fraction of chlorine [START_REF] Knudsen | On the standard-water used in the hydrographical research until July 1903[END_REF]Forch et al., 1902;Barnes, 1954). [START_REF] Knudsen | On the standard-water used in the hydrographical research until July 1903[END_REF] deĄned empirically the relation between chlorinity and salinity S:

S = 0.03 + 1.805 * Cl (1.4)
However, it is not possible to measure the salinity S of a sample analytically and directly, whereas this can be achieved for density and/or chlorinity (Millero et al., 2008). Decades after the Ąrst deĄnition and calibration of the chlorinity standard, a new "Primary Standard" was established in 1937 due to a depletion of the former (Jacobsen and Knudsen, 1940;Culkin and Smed, 1979). This chlorinity standard has been calibrated against silver and was distributed in ampoules for laboratories (Culkin and Smed, 1979). Afterwards, additional salinity deĄnitions have been introduced to express the amount of dissolved salts in seawater while assuming a nearly constant chemical composition of the oceans [START_REF] Knudsen | On the standard-water used in the hydrographical research until July 1903[END_REF].

For example, the electrical conductivity of seawater has been frequently used since the 1950s to account for both, chlorinity and salinity because a proportional relation of conductivity to chlorinity was found (R. Cox et al., 1967). It was deĄned as the conductivity ratio R 15 of the measured electrical conductivity of a water sample to that of water with a salinity of exactly 35 (Wooster et al., 1969) Seawater (EOS-80;Fofonoff, 1985).

The deĄnition of S P is based on "the conductivity ratio K 15 which is the electrical conductivity of the sample at temperature t 68 = 15 °C and pressure equal to one standard atmosphere (p = 0 dbar and Absolute Pressure P equal to 101 325 Pa), "divided by the conductivity of a standard potassium chloride (KCl) solution at the same temperature and pressure" (IOC, 2010). Importantly, and sometimes ignored, the Practical Salinity is unit-less, with S P = 35 at K 15 = 1 (IOC, 2010).

A new standard has been proposed called "Reference Salinity S R " that "represents accurately the Absolute Salinity S A of an artiĄcial seawater solution" (Millero et al., 2008). It is related to Practical Salinity S and can be seen "as an SI-based extension of Practical Salinity, as a reference for natural seawater composition anomalies", and as the currently best estimate for Absolute Salinity of IAPSO Standard Seawater (Millero et al., 2008). The conversion between S and Chlorinity

Cl is S = 1.80655*Cl [g/kg] for 2 < S < 42 (Millero et al., 2008). Further, the authors found for S R the relations to both S and Cl to be: Common pitfalls occurring in both, older and recent literature, by expressing non-uniformly the term "salinity". This is due to the various deĄnitions described above, and the applied method to analyse water samples. For example, soil salinity results in the southern Aral Sea region have been reported consistently in dS/cm because the electrical conductivity was measured (Devkota et al., 2022). On the contrary, Aral Sea water samples were reported in g/kg and parts per thousand (ppt) in the same paper [START_REF] Zavialov | Five years of Ąeld hydrographic research in the Large Aral Sea (2002Ű2006)[END_REF], and in g/L in an older study [START_REF] Létolle | Salinity of surface waters in the Aral sea region[END_REF]. Another method is to analyse a water sample for its total dissolved solids (TDS; commonly in mg/L), a widespread salinity term that has been used to categorise the saltiness of large lakes of the world (Herdendorf, 1990): fresh (<2000 mg/L), brackish (2000-20000 mg/L), saline (20000-40000 mg/L), and hypersaline (>40000 mg/L). An other classical salinity unit is ppt, but also units in mol/kg, g/L or in weight percentage of NaCl (wt.% NaCl, or wt. NaCl) can still be found. Further, it is essential to pay attention to what the concentration of dissolved salts refers to:

• S R ≈ (
• uniquely the volume of solvent in L or mL

• mass volume of the total solvent (solutes includes) in g or kg NaCl (Bodnar, 1993). To present Ćuid inclusions data, the % scale can also be converted into the scale to increase the informative value of data (Bigi et al., 2022). It is important to note that the Šweight % NaClŠ deĄnition refers to the concentration of the ternary system Na-Cl-H 2 O, while it does not refer to seawater salinity scales that had been adapted for complex solutions containing multiple ions.

This section aimed to introduce to the various salinity deĄnitions that exist in the literature. The physical-chemical properties of conductivity, temperature, and salinity are strongly linked with each other, and might be used independently to get rough estimates about the two other properties (Figure 1.4). Therefore, the close relation between temperature and salinity, e.g. in oceans, provides the framework for using two proxy methods of geological records.

As previously mentioned, reported salinity values of water bodies and Ćuid inclusions are not on the same scale, and should always be treated with caution. The relative relationship between individual data in a sample or between sub-samples appears more important here than presenting absolute or exact values that would represent the exact concentration of TDS or Cl. Hence, the use of the same method during a geological salinity study would not affect the data validity in terms of relative salinity changes along a geological proĄle, a core section, a bedded interval or a single crystal.

Chapter 1. Introduction Israel and Jordan (central part: 31°30Š 0Ť N, 35°30Š 0Ť E). The regional climate is inĆuenced by the Mediterranean Sea, causing mild winters and hot summers.

The Dead Sea basin has been occupied by three lakes in its younger history during the Pleistocene and the current Holocene (Lake Amora, Lake Lisan, Dead Sea) [START_REF] Stein | The limnological history of late Pleistocene-Holocene water bodies in the Dead Sea basin[END_REF][START_REF] Stein | The evolution of Neogene-Quaternary water-bodies in the Dead Sea rift valley[END_REF]. Since the DS is a terminal lake and only has the Jordan river as its major freshwater supply, the water balance between inĆow vs.

evaporation/water loss is the dominating factor in terms of the lake-level. The DS is well-known for its rapid lake-level drop since the 1960Šs due to an intensiĄed industrial usage of its water resources and stronger surface water evaporation rates, which have caused a drop of around 30 m from -390 to currently (2020) -435 m below mean sea level with an annual mean lake-level drop of 1 m/year. Therefore, the salinity increased dramatically to ca 34.2 % and former uniform lake has been separated into two basins: the ca 320 m deep northern basin (area ca 45x16 km), and a smaller one in the southern part where salt works are actively precipitating potash for industrial purposes.

The limnological system is sensitive to climate changes and varied over time, e.g. lake-levels declined during warmer periods of Heinrich stadials or increased summer insolation, whereas lake-level rose during colder periods like the Last Glacial Maximum [START_REF] Stein | Abrupt aridities and salt deposition in the post-glacial Dead Sea and their North Atlantic connection[END_REF]Torfstein et al., 2013). However, the lake-level has (probably) never been as low as today, and there are serious concerns about the future water budget that will supply the unstable region with water. However, models showed that the DS will not completely desiccate because a steady state might be reached in a century or later [START_REF] Yechieli | Will the Dead Sea die?[END_REF].

Annual winter period lasts from October to March, the summer period from April-September. Modern-day air temperatures over the Dead Sea are seasonal-controlled and vary accordingly to winter (lowest) and summer (highest), ranging between 18.3 °C (average January) to 32.9 °C (average August) during 1992-2002 (Gertman and Hecht, 2002). Sea-surface temperatures are slightly higher, but follow the same seasonal-controlled pattern with lowest averaged temperatures in January (21.6

°C), and highest in August (33.7 °C). Further, diurnal changes (daily minimum vs. maximum) in air temperatures were monitored during 1992-2002 and reported to vary in winter by 2.7 °C and by 5.2 °C in summer (Gertman and Hecht, 2002). A suitable feature of the Dead Sea basin for palaeoclimate research is a well-preserved sedimentary record of the last 220 ka, documented by characteristic lithologies of aragonite, gypsum, mud, halite and detrital material (Neugebauer et al., 2014).

Hence, speciĄc periods in its younger history (Holocene-Upper Pleistocene) can be investigated by different analytical methods and sedimentological analyses since elevated terraces around the lakeŠs shoreline also bear witness of past climate conditions (Lowenstein et al., 2021;Ghazleh and Kempe, 2021;Lazar et al., 2014).

Furthermore, modern-day observations of in situ halite precipitation revealed a seasonal mechanism that drives coarse halite bottom-growth during summer, and cumulate halite nucleation in the free water body during winter due to a high supersaturation with respect to NaCl, leading to the formation of evaporitic varves (Sirota et al., 2017;Ben Dor et al., 2019). These Ąndings might be an analogue for ancient halite deposits that were described in a long sedimentary record covering several glacial-interglacial periods up to 220 ka ago (Goldstein et al., 2020;see Chapter 5).

Chapter 2. Sampling locations

Aral Sea

The Aral Sea region was a former endorheic large lake, located in Central Asia between Uzbekistan and Kazakhstan (Figure 2.1). It is a further example of anthropogenic exploitation of natural ecosystems and the environmental consequences:

The lake volume has decreased steadily by ca 1000 km 3 since 1960 at varying average rates between 1960-2001 (22.92 km 3 /a), a maximum in 1974Ű1975 (47.59 km 3 /a) and a lowered average rate of 2 km 3 in 2009-2018 [START_REF] Yang | Changes in area and water volume of the Aral Sea in the arid Central Asia over the period of 1960Ű2018 and their causes[END_REF].

Hence, the lake area has decreased as well by ca 60156 km 2 , which has caused a separation into two basins in 1988 [START_REF] Yang | Changes in area and water volume of the Aral Sea in the arid Central Asia over the period of 1960Ű2018 and their causes[END_REF][START_REF] Zavialov | Five years of Ąeld hydrographic research in the Large Aral Sea (2002Ű2006)[END_REF]) (Figure 2.3 D). Since then, the basins were termed ŠSmall SeaŠ and ŠLarge SeaŠ. The Large Sea has been further reduced in its area until the onset of two sub-basins (ŠWestern basinŠ; ŠEastern basinŠ), connected by a small channel (ŠStraitŠ). The Western basin is much deeper (maximum: 43.5 m vs. 7 m), but the salinity increases during the past decades are similar for both basins from 10 to 92 g/kg and 10 to 110 g/kg for the Western and Eastern basin, respectively [START_REF] Zavialov | Five years of Ąeld hydrographic research in the Large Aral Sea (2002Ű2006)[END_REF]. For comparison, the lake was brackish with a salinity ≤12 g/kg before 1960 [START_REF] Zavialov | Five years of Ąeld hydrographic research in the Large Aral Sea (2002Ű2006)[END_REF].

The Aral Sea region resembles the Levante region (including the Dead Sea) in terms of regional aridity, terminal lake state, and rapid decline of lake-level, triggered in both cases by excessive consumption of water resources for industrial and agricultural purposes. Like in the Dead Sea, the negative water balance of the Aral Sea region is accompanied by increasing salinity of the surface and bottom waters [START_REF] Létolle | Salinity of surface waters in the Aral sea region[END_REF][START_REF] Zavialov | Five years of Ąeld hydrographic research in the Large Aral Sea (2002Ű2006)[END_REF]. This caused an intensive salinization of soils and a highly salinity increase of the two main rivers (Amu Darya, Syr Darya) that have been supplied the lake with freshwater from the Central Asian mountain belt [START_REF] Létolle | Salinity of surface waters in the Aral sea region[END_REF]. Not even the glacier melt in the mountain belts due to climate change could equalize the negative water balance of the Aral Sea region [START_REF] Yang | Changes in area and water volume of the Aral Sea in the arid Central Asia over the period of 1960Ű2018 and their causes[END_REF].

The state of the Aral Sea region as a climate archive has been reported in several studies dealing with Holocene lithological units (Oberhänsli et al., 2011;[START_REF] Sorrel | Climate variability in the Aral Sea basin (Central Asia) during the late Holocene based on vegetation changes[END_REF]Sorrel et al., 2006). The main key to interpret different lithologies was found in past lake levels, varying in time according to the amount of water inĆow from the main rivers. A high water inĆow was related with higher runoff of detrital material into the lake, leading to Ąne laminated sediments of clay and salt which hindered the deposition of gypsum and carbonates (Oberhänsli et al., 2011). A reduced water inĆow, in contrast, triggered the precipitation of gypsum and carbonates. The lack of salt units was due to the low salinity and a relatively high SO 4 2 Ű content [START_REF] Zavialov | Five years of Ąeld hydrographic research in the Large Aral Sea (2002Ű2006)[END_REF]. Chapter 2. Sampling locations

Sabkha El Melah

The chosen study site is located in the sabkha ŠEl MelahŠ (SE Tunisia) which is a large depression at the northern edge of the ŚJeffara coastal plainŠ with an area of ca 150 km 2 (Perthuisot et al., 1990;Ben Ameur et al., 2022). The sedimentary record of the sabkha covers the last 5300 years (Fontes and Perthuisot, 1973). The lithology consists mainly of lacustrine and continental deposits, oolithic/biogenic sands, gypsum, huntite, and salt (Davaud et al., 1996;Perthuisot et al., 1990) ( Chapter 2. Sampling locations Gypsum crystals were observed along the shoreline as second evaporitic mineral after carbonates and prior or concurrently mixed with halite. Common gypsum shapes are characterised by needles, Ćat crystals, or lenses, with displacement features within the coastal sediments. Further, larger swallow-tail crystals form Šgypsum beachesŠ (Davaud et al., 1996). Chapter 2. Sampling locations

Archives from the Mediterranean

The accommodation of huge salt and gypsum units in the Mediterranean realm over a short time period (ca 640 ka) occurred during the Miocene and is referred to the ŠMessinian Salinity CrisisŠ (MSC) (Decima and Wezel, 1971;Hsu et al., 1973;Krijgsman et al., 1999). The Ąrst theories that might have caused the rapid evaporitic deposition over a large area were attributed to a desiccated Mediterranean Sea after the closure of the Strait of Gibraltar that interrupted the Atlantic seawater inĆow into the Mediterranean (Hsu et al., 1973).

However, during the last two decades, the MSC has been subject to various studies and debates about key aspects, e.g., the spatial-temporal chronology between locations, the timing of drawdown phases, the water depth formation of evaporites, and models that were designed to link different onshore and offshore sites from Sorbas to Sicily to the Eastern Mediterranean (Rouchy and Caruso, 2006;Roveri, Flecker, et al., 2014). A general consensus was found for the onset of the MSC (currently: 5.97 Ma, based on gypsum cycles in the Sorbas basin, Spain (Manzi et al., 2013)), the division into three internal stages, termed as ŠPrimary Lower GypsumŠ (1. stage), ŠResedimented Lower GypsumŠ (2. stage), and ŠUpper GypsumŠ (3. stage) (Roveri et al., 2008;Manzi et al., 2012), and that the lithological units are related to precessional-controlled cycles of 21 ka [START_REF] Hilgen | Evaluation of the astronomically calibrated time scale for the late Pliocene and earliest Pleistocene[END_REF]Hilgen et al., 1995).

Furthermore, seasonal-derived signatures were discovered in halite and gypsum units in the 2. stage which has been referred to lunar and solar periodicities (Manzi et al., 2012). The end of the Messinian stage is deĄned by the Zanclean ŠGlobal Boundary Stratotype Section and PointŠ (GSSP) at ŠEraclea MinoaŠ section, Sicily, which is also related to the Ćooding of the Mediterranean realm, 5.33 Ma ago [START_REF] Van Couvering | The base of the Zanclean Stage and of the Pliocene Series[END_REF].

The current idea about the formation of the various evaporitic units is mainly based on a model with regional sub-basins in the Mediterranean, each one characterised by differences in water depth, freshwater support, and timing of evaporite deposits (Rouchy and Caruso, 2006;Rigaudier et al., 2011) (Figure 2.6 A-1). There are also approaches that take into account tectonic uplifting in combination with precessional cycles to explain non-marine deposits in a non-desiccated basin (Roveri and Manzi, 2006), and Ša deep-water deep-basin modelŠ that hypothesised the Ćow of dense saline waters towards deeper water depths to be possible causes for supersaturated bottom brines for the formation of evaporites (Roveri, Manzi, et al., 2014).

Chapter 2. Sampling locations

SpeciĄcally the widely distributed gypsum units, both onshore and offshore, have led to studies about the initial environmental conditions during gypsum formation such as salinity, the water depth, and the source of ions by using Ćuid inclusions, petrographical, geochemical and sedimentological analyses (Natalicchio et al., 2014;Natalicchio et al., 2021;Bigi et al., 2022). Experimental results obtained during a preceding thesis have shown that both laboratory-grown and natural halite samples could be used for accurate temperature measurements of the host crystal by combining Ćuid inclusions and Brillouin spectroscopy (Guillerm, 2019). In addition to the Ąnding that it can be used to accurately reconstruct aquatic conditions in halite-bearing terrestrial climate archives, it was further observed that both lake level and volume of the Dead Sea basin (DSB) could be approximately calculated using both, the speed of sound in Ćuid inclusions and the modeled composition of the lake water body at time of halite accumulation (Guillerm, 2019).

Based to these great Ąndings, a major part of the present dissertation was devoted to the extended practical application of BS on halite layers from the sediment core 5017-1 from the DSB. The sediment core 5017-1 has the advantage of having a full record of the climate conditions covering the last 250 ka (Neugebauer et al., 2014;Torfstein et al., 2015; see chapter 2.2.1). Halite layers from the DSB have been interpreted as annual records of drier and/or lower lake levels (Kiro et al., 2016;Ben Dor et al., 2019), thus providing the unique opportunity to reveal palaeotemperatures on annual scales using BS. Hence, two core sections were selected to measure entrapment temperatures of Ćuid inclusions, and to reconstruct the annual evolution during two interglacial periods, the MIS 5e (chapter 5), and the MIS 7c (chapter 6). While halite intervals from core section 5017-1-A-122-2 (MIS 5e) cover a short period of 4-5 years, samples from core section 5017-1-A-176-2

were deposited during more than one decade as successive halite layers on annual scales during MIS 7c.

The aim to study extensively such halite layers is related to recent Ąndings of modern halite depositional mechanisms in the Dead Sea, which showed that seasonal-driven temperature changes have a major control on the formation of two Chapter 3. Material halite types (Sirota et al., 2017):

• cumulate halite (grain size 50-150 µm)

• coarse halite (grain size 400-1000 µm)

It has been observed that cumulate halite precipitates in the mixed water column during winter due to an increased supersaturation of NaCl, while coarse halite forms predominantly during summer when the lake returns to stratiĄed conditions (Sirota et al., 2017). These modern observations are thus an unique opportunity to test qualitatively if such seasonal-driven temperatures changes could be found in the sediment record, and to evaluate if modern observations might act as an analogue for paleohydrological mechanisms in the DSB. [START_REF] Sorrel | The Aral Sea: a palaeoclimate archive[END_REF].

Laboratory-grown gypsum

Chapter 3. Material

Western basin shoreline.

Four water samples could be collected from the Western basin, two of them from small surface ponds fed by upwelling groundwater Ćows (Figure 3.4). Salt crusts around the bottles were observed when the samples arrived in Lyon, and indeed the water differed in its optical appearance from pure water. Both samples were used for sound velocity measurements and labelled as AS11 and AS13. Further, two samples from the Large lake were collected at the surface and at the bottom of ca 20 m depth. Conductivity measurements in 2002 showed that salinity varied from the surface (82 g/kg) to the 25 m lake bottom (110 g/kg) (Friedrich and Oberhänsli, 2004). Hence, a relatively apparent difference in sound velocity can be hypothesised for both water body samples.

Sabkha El Melah

A large (6 cm size) swallowtail gypsum crystal from the Holocene sabkha El Melah was used for detailed studies on trapped Ćuid inclusions using microthermometry and Brillouin spectroscopy to measure Ąnal ice melting points of biphasic FIs and sound velocities of both mono-and biphasic FIs. Its depositional origin has been attributed to the desiccated shoreline of the sabkha in which sulphate-rich waters form bottom-grown gypsum (selenite type).

The growth direction runs in direction to the dihedral form, so that the Ćat crystal bottom (uppermost part in Chapter 4

Methods

Experiments performed during this PhD thesis have been related largely to Brillouin spectroscopy (BS) and included a variety of sample types, technical parameters, analytical approaches, and scientiĄc objectives. Therefore, this chapter aims to provide any necessary information about the experiments, however, all subsequent sub-chapters dealing with BS refer to the methodology described in 4.1.

This chapter outlines several approaches to Ąnally measure:

(1-A) single sound velocity (w) of Ćuid inclusions (FIs) in natural gypsum (4. 2)

and

(1-B) the salinity of FIs using microthermometry (4.7),

(2) multiple w to calculate entrapment temperatures (T x ) of gypsum and halite (4.3), and

(3) single w of capillary solutions to build a salinity proxy based on BS (4.6).

Brillouin spectroscopy

This section provides technical information about the basic principles of Brillouin spectroscopy and the instrumental set-up using aligned optics in the laboratory. For information about the practical procedure during measurements of Ćuid inclusions, the reader is referred to chapters 5.3 and 5.7.

Principles

The Brillouin spectroscopy method (BS) is named after the French physicist Léon Brillouin (1889 -1969) who predicted theoretically for the Ąrst time how x-rays would propagate through crystals by inelastic scattering of light (Brillouin, 1922).

The method is based on scattering effect due to an inelastic interaction of an incoming monochromatic light and thermodynamic Ćuctuations of any substance.

Travelling light is characterised with a frequency f and corresponding to photons of energy:

E = h * f (4.1)
with E the energy and h the Planck constant. The light-matter interaction can be either elastic (no energy change) or inelastic (energy increase or decrease). The principle of BS is the inelastic scattering of light, which induces a change in energy of the photons, associated to a frequency shift, the Brillouin frequency shift ∆f B ).

The main principle for BS measurements is the relationship between ∆f B and sound velocity w at a given laser wavelength λ and refractive index n, respectively:

∆f B = 2 * sin( θ 2 ) * n * w λ (4.2)
where θ is the scattering angle. In our case, backscattering is used in the microscope objective, which corresponds to θ = 180°.

Optical setup

The spectroscopy laboratory in Lyon (Institute of Light and Matter, Campus de la Doua) is equipped with a Verdi V-6 system (Coherent) that is a diode-pumped Nd:Vanadate (Nd:YVO 4 ) laser, working with green monochromatic light (532 nm).

To cool the system, it is supplied externally by both an air-conditioning system and a chiller that is connected with water lines to the system. The common working laser beam power is set to 4 W. This high power is required to ensure

m(Q, t) = (c 2 ∞ -c 2 ) * Q 2 * [exp -( c c ∞ ) 2 * t τ ] (4.3)
where Q is the wavevector probed by the scattering setup, c and c ∞ the sound velocity at zero and inĄnite frequency, respectively, and τ a structural relaxation time. Note that c ∞ is chosen to be a constant at 3000 m/s [START_REF] Pallares | Anomalies in bulk supercooled water at negative pressure[END_REF]. Further, Q is given in backscattering geometry by:

Q = 4 * π * n λ (4.4)
where n is the refractive index of the liquid. The full analysis includes further both the instrumental resolution function (IRF) and the dark count of the setup.

The IRF was determined for each pinhole (entrance of Brillouin box in Figure

Sound velocity measurements

An individual sound velocity measurement of FI i follows a speciĄc protocol that has to be respected with respect to its workĆow, security standards, and data organisation (see Appendice A). Suitable locations of FIs were initially investigated using objectives x10, x20, x50 or x100.

Gypsum samples were mostly measured at seven to eight, partly even at more than eight temperature points. The Ąrst experiments were initially carried out at room temperatures because (1) the Linkam stage does not need to be cooled using liquid nitrogen (N 2 ), ( 2) the potential risk to nucleate a vapor bubble is prevented, and

(3) potential damages during sample preparation can be detected by analysing the Ąrst batch of acquired spectra. Regardless of the FI phase, the most secure way was found to start at 22 °C < T < 25 °C. Subsequently, the crystal (including all FIs) was cooled down gradually by maintaining a distinct temperature difference between each temperature step, e.g., a 8 °C difference would be respected between the starting measurement I (23 °C), the followed measurements II (15 °C) and III (7 °C), and the last measurement IV (31 °C). These thermal cycles can be adapted for each sample with varied temperature differences.

It is important to note that the lowest temperatures (III) should be well < 10 °C since local thermal Ćuctuations in natural brines can occur. However, for natural samples from locations with warmer climates (e.g. sabkhas, deserts, thermal springs), the temperature ranges can be adapted towards higher values. As an example, FIs in gypsum samples from Tunisia with known brine temperatures well > 20 °C were measured for Brillouin Spectroscopy at elevated temperatures between 15-50 °C to investigate their potential entrapment temperatures.

To analyse a dataset of crystal i with n sets of FI i s, a registered .txt-Ąle after acquiring the frequency shift has to be analysed with a special analytical code using Wolfram Mathematica. This .txt-Ąle contains all essential information about

Entrapment temperatures

The procedure for entrapment temperature measurements in FIs is slightly more advanced than for single w measurements : Although the main principle of acquiring ∆f B spectra does not change, it is based on six to eight runs per FI. E.g., for halite the most accurate way to produce a signiĄcant amount of FI data are eight single measurements (Guillerm, 2019;see chapter 5). The choice of intervals to be selected is left up to the sample type, the mineral, or the scientiĄc objective, however, it should be well thought out before starting the work.

Entrapment temperatures (T f ) are deĄned as the intersecting point of Ątted isochores and SLVE curves at ambient pressure. Hence, single w values obtained in 4.2 had to be sorted by descending temperature values and for each FI per crystal in separated Ąles for isochores and SLVE data, allowing to Ąnd each intersection by calculating the crossing point (T x ) between a line (isochore) and a polynomial (SLVE). The calculation of T x was carried out by using a program code in Wolfram Mathematica. A monophasic FI follows a quasi-isochore (dashed black line) upon heating (red arrow) with increasing internal pressure. During cooling (blue arrow) the internal pressure becomes negative. The FI stretches to to overcome the metastable phase (dashed black rectangles). Energy release forces a vapor to nucleate to re-enter the stable phase. By theory, the thermal path of a biphasic FI follows SLVE that represents ambient pressure conditions. Right: Sound velocity measurement in monophasic (black dashed line; isochore) and biphasic (red solid line; SLVE) phases by at four temperature steps. The intersection of both lines (T x ) gives the entrapment temperature (T f ) if the sample has precipitated at ambient pressure. In most geological cases, a pressure-correction is needed.

Thermal cycles of Ćuid inclusions

The analysis of entrapment temperatures in FIs of any kind of mineral requires two thermodynamic phases as shown in Section 4.1. The measurement of monophasic FIs is carried out simply by varying the temperature in the Linkam Stage. For a moderate temperature interval (e.g., between 5-40 °C), the resulting pseudoisochoric curves for w(T) are linear. In general, w(T) along the solid-vaporequilibrium (SLVE) is not linear, and we use a polynomial Ątting to describe the SLVE. An issue is that it is not always straightforward to nucleate a bubble in a monophasic FI to make it biphasic. A usual method is to sufficiently cool the sample in order to induce a tension in the liquid high enough to trigger cavitation.

Unfortunately, this does not happen systematically. Experiments carried out on both synthetical and natural halite (NaCl) samples showed that salt-rich FIs cavitated rapidly (within seconds to minutes) after thermal heating-cooling cycles from ambient temperatures up to 130 °C and back within 2-4 minutes (Guillerm, 2019). Heating results in permanent plastic deformation of the FI, with a volume increase that makes the tension higher when cooling back to ambient temperature, leading to cavitation. This procedure was found to transform FIs reliably from monophasic into biphasic phases in most cases, enabling to perform measurements Chapter 4. Methods of both phases within the same day of experiments (Guillerm, 2019).

However, thermal-heating cycles might be limited for the four phases of the [START_REF] Hoff | Untersuchungen über die bildungsverhältnisse der ozeanischen salzablagerungen[END_REF]. This is due to the low dehydration temperature at atmospheric pressure that starts slowly above 42 °C, with a metastable Ąeld between 42-97 °C (Posnjak, 1938). Other studies showed that dehydration to hemi-hydrate starts >80 °C and further >115 °C to transform into the anhydrous phase, depending on the external partial pressure of water vapor (Ball and Norwood, 1969). Nevertheless, two kinds of thermal cycles have been tested to Ąnd the best cavitation procedure for gypsum samples: (1) heating-cooling cycles (hcc) and ( 2) cooling-heating cycles (chc) at several temperature steps, temperature ranges and heating or cooling rates. The stepwise heating during hcc should test if gypsum FIs cavitated faster and more often compared to a direct heating from room temperature to +130 °C. During chc, more negative temperature steps were tested (e.g., -20, -40, -60, -80, -100 °C) with varying duration. Temperature ranges of hcc (chc) did not extend beyond 130 °C (-130 °C). The rates of heating or cooling were either set to be slow (e.g., 2-10 °C/min) or fast (e.g., 10-30 °C/min) to study the control on gypsum FI cavitation processes with respect to the thermal rates.

CaSO 4 Ű H 2 O-system with gypsum CaSO 4 • 2 H 2 O, hemihydrate CaSO 4 • 1 2 H 2 O, y-CaSO 4 and anhydrite

Salinity

As introduced in chapter 1.5, The term "salinity" refers to a variety of used deĄnitions and units. For this work, both published data and derived equations for NaCl solutions were used for interpreting the measured sound velocity data (see Chapter 8.1,4.1). This allowed to compare and evaluate directly the data from the literature with measured sound velocity values of pure water and NaCl solutions.

Following this, other solutions containing natural brines from the Mediterranean Sea, the Aral Sea and the Dead Sea were used to estimate the salinity values.

A crucial point in interpreting sound velocity data from Brillouin spectroscopy is the correction factor for the capillaries because the refraction of light at interfaces affects the detected Ąnal w value. The correction factor is determined by measuring the raw value of w in a capillary Ąlled with pure water. The average ratio of the tabulated values for w to the raw values is 0.9741 ±0.002 (8.1).

The temperature to account for the correction factor evaluation was set to T = 25 °C because it turned out to be the most practical value during BS-salinity measurements (see chapter results and discussion). The relative NaCl w values from Chen and Millero, 1986 are converted into absolute ones, using 4.5: w 0 H2O = 1402.388 + 5.03711T -5.80852 * 10 -2 T 2 + 3.342 * 10 -4 T 3 -1.478 * 10 -6 T 4 + 3.1464 * 10 -9 T 5 (4.5)

with T the temperature in °C. After 4.5, w 0 H2O (25 °C) = 1496.7 m/s, which is used as constant summand for absolute NaCl w values. Both the relative and absolute NaCl w values from Chen and Millero, 1986 are listed in Table 4.1. The derived equation from the cubic Ąt is:

w 0 N aCl = 0.031m 3 -2.3728m 2 + 61.82m + 1497.2 (4.6)
with m the concentration (molar ionic strength) in mol/kg or mol/L. The known concentrations of NaCl-bearing capillaries have been inserted into 4.6 and compared to data from the literature (Table 4.1).

The observed Ąnal ice melting points (in °C; see 4.7) were used to calculate the salinity S, expressed here as "weight % NaCl equivalent" (wt. % NaCl eq.). The salinity of each FI was calculated after Bodnar, 1993:

S = -1.78 * T ice * (-1) -0.0442 * T 2 ice + 0.0005572 * T 3 ice * (-1) (4.

Saline aqueous solutions

First, the mass of NaCl to be weighed in x mass water (m H 2 O ) for a certain desired salinity S i (in wt %) was calculated by:

m N aCl = - m H 2 O * S i S i -100 (4.8)
with m N aCl in g, m H 2 O in g, and S i in %. For a molar mass M of NaCl of 58.44 g/mol, the molality b (in mol/kg) was calculated using:

b = m N aCl * 1000 M * m H 2 O (4.9)
Under air-conditioning controlled conditions at 21 °C, the masses of NaCl (UMR Chemicals NORMAPUR, 100% NaCl) of S i were Ąrst weighted dry into 100 mL beakers using a balance (Sartorius CPA225D, ±0.01 mg), and then Ąlled up to 50 g with pure H 2 O (Milli-Q® 18.2 MΩ-cm at 25 °C). The sample was then placed on a magnetic stirrer (Cat MCS66) and homogenised using a magnetic stirrer bar until the salt was completely dissolved (ca 30 minutes of stirring After sealing, speed of sound measurements were started without delay at 5 °C until maximal 65 °C. A typical Brillouin spectra needs to reach at least 300 counts during its acquisition so that the frequency shift can be analysed precisely Ąnally.

Samples were acquired three times at each temperature step (≈ 5 seconds per acquisition). The Brillouin spectra were analyzed by Ątting the acquired frequency shifts in the same package of Wolfram Mathematica as used for FIs.

A list of all NaCl mass solutes, molalities, and salinities is given in Table 4.2.

Final ice melting temperatures

The essential step to derive the salinity of a sample using microthermometry on Ćuid inclusions, was to freeze the sample and to cavitate a vapor bubble to avoid metastable conditions (Roedder, 1984; see chapter 4.5). Then, the temperature of the Linkam stage was slowly increased from negative low temperatures, and the point, at which the last ice crystal within a FI melts, was documented and labelled the Ąnal ice melting point of the studied FI.

A natural gypsum sample from a sabkha, Tunisia, was prepared using a razor blade to cut small pieces (ca 1.5 cm) along the 010 plane. The pieces were placed in a temperature-controlled stage (Linkam THMS 600) with 0.1 °C resolution (Figure 3.5). To Ąnd out whether both sound velocity and Ąnal ice melting temperatures differ signiĄcantly between two more distant parts of the same parent sample, >30 growth bands with primary FIs were selected for BS analysis (Figure 3.6 A).

Petrographic observations of FIs were carried out using reĆected light with an objective of magniĄcation x100 of the same microscope as described in chapter 4.1. The selected FIs were documented with a camera attached to the microscope.

Post-depositional processes affecting the pristine signal of FIs (e.g. by external forcing such as burial, exposure, dissolution-(re)precipitation, necking down, etc.) have been ruled out by selecting only primary monophasic FIs that are aligned along straight growth bands.

The samples were cooled from room temperature at a rate of ca 20 °C/min to -96 °C for 3-5 minutes to ensure a completely freezing of all investigated FIs. Subsequently, the sample was heated again at ca 10 °C/min. The rates were reduced to 1 °C/min when phase changes from liquid+vapor to solid+liquid+vapor were expected to occur. To determine if the chemical system of the FIs would be H2O-NaCl, the eutectic point was documented when FIs started to melt. For H2O-NaCl the eutectic point was determined to be at -21.2 ± 0.1 °C [START_REF] Sterner | Synthetic Ćuid inclusions in natural quartz I. Compositional types synthesized and applications to experimental geochemistry[END_REF].

Above this mark (i.e. > -21.2 °C), the rate was reduced to about 0.5 °C/min to be able to document the Ąnal ice melting point. Care was always taken to ensure the presence of a gas bubble, as otherwise the exact determination of salinity cannot be guaranteed (Roedder, 1967). He observed ice-liquid equilibrium in pure water at 6.5 °C, because of tension. The observed data were then grouped by sub-sample and growth band and plotted against sound velocity (see Section 4.2). were entrapped in the halite crystals which serve as windows to estimate the chemistry and physical properties of the primary lake water conditions. Brillouin et al., 2000;Denton et al., 2005). A substantial impact of seasonality on various proxies has been proposed (Clemens and Prell, 2003;Huguet et al., 2006) at annual (Brocas et al., 2018;Felis et al., 2004), centennial (Bar- Matthews et al., 2000;Brauer and Dulski, 2009) or millennial-scale resolutions (Milner et al., 2012).

Continental archives such as tree rings, laminated speleothems and varved lake sediments can provide seasonal records of different climate variables (Shi et al., 2010;Baldini et al., 2021;Ben Dor et al., 2019). For example, laminations in speleothems commonly accumulate more calcite in the cool season and less or none during the warm season (Riechelmann et al., 2020). Analogously, varved sediments have been extensively used for the last few decades, to reconstruct, e.g., seasonal rainfall intensity in temperate/periglacial lakes (Ariztegui et al., 2010), the long-term progressive lake-level decline in a hyper-arid region in northern Chad (Francus et al., 2013), both winter precipitation and temperatures in southern Finland (Ojala et al., 2013), Ćooding events in the Southern Alps [START_REF] Wirth | A 2000 year long seasonal record of Ćoods in the southern European Alps[END_REF] as well as Holocene temperatures in China and Switzerland (Chu et al., 2005;Trachsel et al., 2012).

Calibration studies using the δ 18 O composition of modern ostracods in turn provide temperature reconstructions that are in good agreement with the observed range of seasonal changes in lake water temperature (Labuhn et al., 2021), based on ŞpalaeotemperatureŤ equations for equilibrium carbonate precipitation and the estimated mean δ 18 O of lake water. However, the high intra-species carbonate δ 18 O variability of samples leads to a large scatter in the reconstructed temperatures often compromising palaeotemperature reconstructions. Recent progress in reconstructing palaeotemperatures has been made by using biomarkers (i.e., leaf wax n-alkanes, bacterial branched glycerol dialkyl glycerol tetraethers (brGDGTs)), however, these geochemical proxies include certain limitations as well (Wang et al., 2018;Liu and An, 2020). Thus, most palaeoclimate proxies provide relative changes in climate variables while a direct and precise reconstruction of such variables is still lacking. Therefore, there is a crucial need for resolving seasonal temperature The hydrographic lake regime has been switching between meromictic (stable stratiĄcation without mixing during a year) and holomictic (lake mixing once a year) periods since 1978 (Gertman and Hecht, 2002;Anati et al., 1987). The long-term meromictic regime of the modern Dead Sea was interrupted for just two months with an overturn of the water body in winter 1978Ű1979 (Steinhorn, 1985), associated with the onset of halite precipitation on the surface in February 1979 (Steinhorn, 1983). After several switches between both regimes (meromictic phases from 1979Ű1982 and 1991Ű1995; holomictic phase from 1983Ű1991), perpetual holomictic conditions have prevailed since 1996 coevally with the precipitation of halite (Gertman and Hecht, 2002;Anati et al., 1987;Anati and Stiller, 1991).

Moreover, seasonal-driven variations in solar radiation, temperature, quasi-salinity and the degree of halite supersaturation have been both observed and extensively studied (Arnon et al., 2016;Gertman and Hecht, 2002;Sirota et al., 2017). The main seasonal feature is a well-mixed water body from surface to bottom during the winter season between late November to early March (Gertman and Hecht, 2002), and a thermal stratiĄcation from early spring to late fall (Gertman and Hecht, 2002). Modern halite precipitation is characterised by ( 1) bottom-grown coarse halite (mainly springŰlate summer) and ( 2) Ćoating Ąne cumulate halite that nucleates in the whole water column excessively during winter due to a high supersaturation of halite, and to a minor extent during summer below the thermocline (Sirota et al., 2017).

The DSB has been successively occupied by several lakes and water bodies since the Ąrst marine incursion from the Mediterranean Sea during the Miocene (Neev and Emery, 1967;Stein, 2001). The current study focusses on the sediments from Lake Samra, which was the last interglacial (ca 129Ű116 ka) precursor of the Dead Sea (Waldmann et al., 2009) A recently improved age model for core 5017-1-A (Figure 5.1 D) includes a new tie point at 129 ka ± 1 ka at 316.24 m blf (i.e., close to the bottom core section at 315.17 m blf), which therefore corresponds to the onset of MIS 5e during the last interglacial period (Goldstein et al., 2020).

Methods

Sampling of core section 5017-1-A-122-A-2 was carried out at GFZ Potsdam (Germany). Selected pieces of halite crystals were placed on their Ćat side on a microscope cover slip in a temperature-controlled stage (Linkam THMS 600) with 0.1 °C resolution. Samples were studied with an upright optical microscope (Zeiss Axio Imager.Z2 Vario), equipped with a long-working distance x100 objective (Mitutoyo Plan-Apo, N.A. 0.7). To select suitable FIs for analysis, published criteria for accurate Brillouin analysis on halite crystals (Guillerm et al., 2020) was followed. For palaeothermometry, the previously described full set-up and detailed physical principles of BS were used (Mekki-Azouzi et al., 2015).

The method is based on the measurement of sound velocity (w) in a liquid which depends on the composition, temperature (T ) and pressure of the liquid. When heated or cooled, a monophasic FI (only liquid) follows an isochore at constant volume, along which w varies with T . In a biphasic FI (liquid + vapour), w of the liquid is measured as a function of T at the solid-liquid-vapour equilibrium (SLVE).

Hence, a transformation (cavitation) from a monophasic phase into a biphasic phase is required for each FI and achieved by a short-term thermal cycle (Figure 5.5). By deĄnition, for a FI trapped near ambient pressure (0.1 MPa), the crossing temperature T x equals the entrapment temperature T f at which both w curves intersect (along both the isochore and SLVE; S1 and S2).

A value for one Ćuid inclusion (FI) consists of eight single measurements conducted Each mean T x value is calculated by the mean of all measured FIs:

T x,mean = x i N (5.1)
with x i the T x value for FI i and N the number of all FIs. Mean T x values are analysed statistically by means of their standard deviation (SD), derived from equation ( 2):

SD = (x i -T x,mean ) 2 (N -1) (5.2)
The standard error of the mean (SE), used as error bars in each Ągure, is calculated by equation (3):

SE = SD √ N (5.3)
To rule out post-entrapment processes in the samples, histograms were used to check for statistical distributions that are close to Gaussian distributions (Figure

5.7).

A correction term was applied on measured T x data according to the pressure of the water column above where crystals form (T f = T x + ∆T P , with ∆T P = 4.8 °C for crystals precipitated at lake bottom). Here, a water column height of 539 m was calculated based on a pressure of 6.45 MPa and a lake-level of -250 m above sea level (Guillerm, 2019). Such a reconstructed lake-level value, based on speed of sound measurements in biphasic FIs (Guillerm, 2019), is similar to that of a recently published synthetic lake-level for the beginning of the Last Interglacial (Torfstein et al., 2013). It is noteworthy to mention that, as the coarse halite layers Chapter 5. Reconstructing lake bottom water temperatures and their seasonal variability in the Dead Sea Basin during MIS 5e studied here were deposited during a short time interval, the pressure correction is virtually equal for all of them. Therefore, any difference between uncorrected T x can be directly converted into a difference between T f values. Conversely, as cumulate halite crystals form in the water column between the surface and the bottom of the lake the unknown T correction for formation depth is less or equal to the T correction for coarse halite crystals. Therefore, in view of the T x values, the T f of cumulate halite layers must be less than the T f for the surrounding coarse halite layers.

Results

Individual the surrounding coarse halite layers.

Discussion

The studied section of cumulate-coarse alternations provides insights into a short depositional period (ca 3 years) during which Lake Samra was most likely holomictic with a seasonal precipitation mechanism (Kiro et al., 2016). Another type (white, brown and black laminae with Ąne detritus and/or gypsum) of the layered ŠlhŠ facies in core 5017-1-A has already been related to annual deposition cycles with periodicities of ca 11, 7Ű8, and 4Ű5 years (Palchan et al., 2017). However, the sampled section in this study lacks clear detrital layers and only consists of cumulate-coarse alternations. The fact that alternating detrital layers are missing indicates rather dry depositional conditions without Ćood events or increased runoff into Lake Samra during cumulate-coarse halite deposition (Ben Dor et al., 2019). However, the ŠlhŠ facies type described by Palchan et al., 2017 is observed below and above the studied section in core 5017-1-A-122-2 (Figure 5.2 B), indicating a full annually deposited sequence with changing hydrological parameters that drove halite precipitation and clastic deposition. Further, the studied section can be considered a seasonal record of winter (cumulate halite) and summer (coarse halite) as previously observed in the modern Dead Sea (Sirota et al., 2016) and Hecht, 2002;Gertman et al., 2010). The uncertainties are reported as both, standard error of the mean (SE; error bars), and 95% conĄdence intervals (black lines). CO layers highlighted by vertical bands. (C) Monthly monitored deep lake temperatures between 2004Ű2010 (blue squares; from Gertman et al., 2010). The vertical bands illustrate the approximate period during which CO halite crystals grow on the deep lake Ćoor.

Chapter 5. Reconstructing lake bottom water temperatures and their seasonal variability in the Dead Sea Basin during MIS 5e

have revealed annual to seasonal temperature variability during MIS 5e. The δ 18 O values and Sr/Ca ratios in corals from the Red Sea reported sea surface temperatures (SST) with a seasonal amplitude of 8.4 °C at 122 ka (Felis et al., 2004), whereas pollen records enabled the reconstruction of seasonal temperature and moisture changes for the last 150 kyr (Chen and Litt, 2018). Based on their deĄnition of pollen assemblage zones (PAZ) for core 5017-1-A, the sampled section at 315.15 m blf falls into the end of PAZ IV2 (340.6 -313.7 m blf). This period marks the start of a transition from humid to dry conditions in the Levantine region by a progressive aridiĄcation, which is in line with the onset of the Ąrst halite deposition at the beginning of MIS 5e (Chen and Litt, 2018). Further, the authors report an initial warming starting in PAZ IV1 (313.7Ű290.2 m blf) judging by the signiĄcant pistachio pollen increase at this time. Since the sampled section is slightly older, it is suggested that the obtained lake water temperatures of cumulate/coarse halite layers might still reĆect relatively colder conditions under an increasing drying trend in the region.

Sedimentary records in Europe covering the MIS 5e interval are more abundant.

For instance, reconstructed mean temperatures of the coldest and warmest months in South-West Europe (Alboran Sea) allowed deĄnition of the SST seasonality to 6 to 7.5 °C on millennial timescales during the MIS 5e, coevally with an increase in precipitation and highly evaporative summer conditions (Martrat et al., 2004).

In southern Italy, a carbonate varved section of Lago Grande di Monticchio (127.20 ka; associated with the onset of MIS 5e) exhibits a rapid increase in Quercus pollen (128.85 ka) and abundant pollen of Mediterranean woody taxa (Brauer et al., 2007). This has been interpreted as intervals of hot summers, seasonal moisture deĄciency and high Ąre frequency, associated with peak summer insolation (Brauer et al., 2007). Alternatively, speleothems from the Soreq Cave reported an increase in annual rainfall on centennial scales at the onset of the MIS 5e, associated with frost-free winters and drought-free summers in the Levantine region (Bar- Matthews et al., 2000). Most of these records, however, are missing a precise quantiĄcation of the interpreted palaeoseasonality.

Annually grown halite layers in the Dead Sea can be used to overcome such limitations by considering modern halite deposits as an analogue for the abundant halite layers in DSDDP core 5017-1-A (Figure 5.2 A). The onset of modern CO halite growth coincides with a seasonal transition, when lake bottom temperatures are comparatively colder during winter (Sirota et al., 2017) Chapter 5. Reconstructing lake bottom water temperatures and their seasonal variability in the Dead Sea Basin during MIS 5e between early and late summer lake bottom temperatures (Gertman et al., 2010;Arnon et al., 2016;Sirota et al., 2016). The results of this study follow this pattern with increasing hypolimnetic temperatures from the base to the top of CO halite layers in DSDDP core 5017-1-A-122-2 (with an increase of 3.0, 4.6 and 1.5 °C for CO1, CO2 and CO3, respectively; Figure 5.3 B). Hence, CO top samples can be interpreted as palaeo-summer halite precipitates. Accordingly, CO bottom samples (i.e., CO1-BOT, CO2-BOT, CO3-BOT) most probably recorded cold (late winter and early spring) hypolimnetic temperatures. The results thus highlight that holomictic periods in Lake Samra must have been associated with coarse halite precipitation in summer and seasonal temperature variations in the hypolimnion.

Further support for large seasonal water temperature differences can be found in cumulate halite samples and their modern analogues: the Dead Sea starts mixing in late NovemberŰearly December today, with both a cooling of the brine and a (nearly coeval) increase in NaCl saturation which drives nucleation of cumulate halite in the homogeneous water column (Sirota et al., 2017). Thus, the lower T x values obtained in cumulate halite FIs can be considered as seasonal winter records in this study. This is further in line with higher palaeotemperatures in both surrounding coarse halite samples CO2-TOP and CO3-BOT, which reĆect the end and the onset of two successive summer seasons, respectively (Figures 5.3 B;5.4). However, it is important to note that the uncorrected cumulate T x values do not picture the factual palaeoseasonal temperature variations with respect to the surrounding coarse layers. Since the initial (absolute) water depth of cumulate halite nucleation remains unknown, only the lower (= T x ; ca 16 °C) and upper boundary (= T x + 4.8 °C; ca 20.8 °C) can be estimated for the studied cumulate layer. Hence, the palaeoseasonality would range between ca 2Ű7 °C which prevents the absolute temperature differences from being assessed quantitatively, e.g., between early spring/deep winter or late summer/deep winter. The focus, therefore, will be on entrapment temperatures obtained from coarse halite samples.

In accordance with this, the averaged T f of coarse-bottom samples (19.4 ± 1.1 °C), which reĆects late winter water/early spring lake bottom temperatures, is lower than the present-day lake bottom temperature in winter of ca 23 °C. On the contrary, the averaged T f of coarse-top samples (22.5 ± 1.1 °C) is close to present-day mean summer maxima (23.8 °C for the period 1996-2010; Figure 5.3 B).

The offset between palaeo-late winter and present-day winter values of ca 4 °C could be explained by an enhanced seasonality during MIS 5e. A lower winter insolation (200 W/m 2 versus present day 220 W/m 2 ) and enhanced summer insolation at both 30°and 35°North at that time may have driven the regional climate towards colder winters over the south and central Levant (Kiro et al., 2017;Torfstein et al., Chapter 5. Reconstructing lake bottom water temperatures and their seasonal variability in the Dead Sea Basin during MIS 5e 2015). This is in line with climate simulations that showed a southward shift of winter storm track rains in the Mediterranean Basin and the Levant, linked with a strengthened Siberian anticyclone that pushed colder air farther south in Europe (Kutzbach et al., 2020). Further, the climate model indicated slightly wetter winters and no change or even drier conditions at times of precession-caused high seasonality (ca 125 ka) in summertime (Kutzbach et al., 2020). Interestingly, a stronger palaeoseasonality was also reported in Red Sea corals during the peak of MIS 5e based on increased sea surface temperatures (with seasonal cycles of 8.4 °C at 122 ka), compared to a modern seasonal amplitude ranging between 4.5 and 5.6

°C (Felis et al., 2004). The weaker modern seasonality in the Levant and eastern Mediterranean Sea region has been associated with enhanced winter insolation and lower summer insolation in the Northern Hemisphere during the Holocene (Kiro et al., 2017).

The modern amplitude of seasonality, measured as the difference between winter and summer lake bottom water temperatures, exhibits a nearly constant value of 1.5 °C for the period 1996-2010 (Figure 5.3 C). This modern seasonal thermal amplitude is thus at least half lower than the reconstructed seasonal amplitude from coarse halite data for the MIS 5e (Figures 5.3;5.4). The results therefore suggest a stronger palaeoseasonality during MIS 5e, with distinct differences in entrapment temperatures between the bottom and the top of each coarse halite layer (Figure 5.4).

Conclusions

Using the speed of sound in halite Ćuid inclusions from the Dead Sea basin, this study reconstructs lake-bottom temperatures of Lake Samra during Marine Isotope Stage (MIS) 5e. As demonstrated by three successive coarse halite layers, palaeotemperatures increased between the bottom (late winter/spring) and the uppermost (late summer) sampling interval of each coarse layer. Conversely, three sub-samples of a Ąne cumulate halite layer yield stable and lower (deep winter) temperature values, highlighting that the alternation of coarse and cumulate halite layers in core 5017-1 reĆect seasonal contrasts of lake-bottom waters under holomictic conditions during MIS 5e. Our data further imply that the reconstructed seasonal amplitude of lake-bottom water temperatures at the onset of MIS 5e in Lake Samra was at least twice as high as the modern amplitude of seasonality.

Hence, this study clearly highlights the potential of using BS to investigate primary FIs in annual salt layers and reconstruct palaeotemperature variations in lacustrine sediment cores, at a seasonal resolution. Moreover, using BS on base-top Chapter 5. Reconstructing lake bottom water temperatures and their seasonal variability in the Dead Sea Basin during MIS 5e halite crystals from multiple halite layers does not require any preparation of thin sections, is easily applied on small FIs, and can provide a reasonable quantity of T x measurements. Thus, future studies dealing with (palaeo) Dead Sea-like hydroclimates should focus on coarse salt layers to reconstruct the seasonal variability of lake temperatures on decadal to millennial scales.

Chapter 5. Reconstructing lake bottom water temperatures and their seasonal variability in the Dead Sea Basin during MIS 5e

Supporting information

This part contains supporting Ągures (S1 -S4) with legends related to the described analytical procedure in the methods.

Introduction

A pioneering step of our understanding in global EarthŠs climate processes has been made by identifying variations of stable isotope compositions derived from Pleistocene-Holocene climate archives (Martinson et al., 1987;Petit et al., 1999).

Such variations were recognised by isotopic excursions and have led to the deĄnition of Marine Isotope Stages (MIS) that correspond to glacial (cold) and interglacial (warm) periods, with synchronous trends in continental ice volumes, and both atmospheric CO 2 concentrations and temperatures (Figure 6.1). Each period is characterised by the presence (glacial) or absence (interglacial) of major continental ice sheets that thus have a major control on global sea-level.

The general MIS numbering is based on marine fossil foraminifera and has provided a detailed palaeoclimatic record of the past 5 million years on millennial to centennial resolution (Lisiecki and Raymo, 2005). For example, MIS 7 is divided into Ąve sub-stages (labelled from a to e), while MIS 7e is the oldest at the boundary to Termination III (ca 246 ka ago), marked by a rapid shift from glacial to interglacial conditions (Tzedakis et al., 2004) (Figure 6.1).

However, terrestrial climate archives are often affected by stratigraphic gaps (hiatus) due to erosion or deposition shortages (omission). Therefore, well-resolved climate records are scarce and not uniformly distributed for all regions of interest. Nevertheless, ice core and terrestrial-derived pollen records have helped to resolve palaeoclimatic periods during Early-Middle Pleistocene MIS sub-stages in, e.g., Antarctica (Petit et al., 1999), Britain (Candy and Schreve, 2007), Portugal (Tzedakis et al., 2004), or Greece (Tzedakis et al., 2003).

The Vostok ice core revealed four prominent peaks in atmospheric CO 2 concentrations during MIS 11.3, 9.3, 7.5 (=MIS 7e), and 5.5 (=MIS 5e) (Petit et al., 1999). Further, records have marked the MIS 7 period as an interesting study interval that might be an analogue to present-day global warming, with similar global sea-levels during 7c and 7e compared to present-day (Tzedakis et al., 2004).

It is further characterised by three warm shorter sub-stages (7e, 7c, 7a) which revealed short periods of strong atmospheric warming in Antarctica which coincides with increased CO 2 concentrations, decreased ice volume, and more negative δ 18 O signatures (Figure 6.1; Watanabe et al., 2003). The site Dome Fuji (E Antarctica) provided a millennial-scale record showing an increase by 8 °C around 220 ka ago during the MIS 7d-7c transition (Watanabe et al., 2003). Support for such an abrupt warming 220 ka ago comes also from marine sediment records retrieved 100 km offshore of NW Spain (42°N), where reconstructed sea-surface temperatures from planktic foraminifera increased by ca 10 °C (summer) and 8 °C (winter) (Desprat et al., 2006). DeĄned MIS stages plotted together with atmospheric oxygen isotope composition (Sowers et al., 1993;Jouzel et al., 1996;Malaizé et al., 1999), ice volume changes, atmospheric CO 2 variations, and Antarctic temperatures reconstructed from oxygen isotopes (Petit et al., 1999). Glacial-interglacial sub-stages of MIS 5 and 7 indicated by a to e, climatic transitions marked by Terminations I-IV. LIG: Last interglacial. Right: MIS 7 with three short-term warming sub-stages (a, c, e) and two glacial intervals (b, d), reconstructed by planktonic foraminiferal oxygen isotopes and tree pollen abundances (Tzedakis et al., 2004), and deuterium isotopes from the Vostok ice core (Petit et al., 1999). Yellow vertical bands in both plots show the halite units studied in this thesis (MIS 7c = 5017-1-A-176-2; MIS 5e = 5017-1-A-122-2).

The Climatic Optimum of MIS 7 has been reported for the Bay of Biscay (North Atlantic; 46°N) between 220-210 ka based on benthic foraminiferal species yielding an oxygen isotope excursion, coinciding with a strong increase in dinocyst species at maximum July insolation at 65°N (Penaud et al., 2008).

However, none of these climate records could achieve a temporal resolution better than centennial scales, hence terrestrial palaeoclimate studies are often temporarily limited. Although orbital processes had the major control on ancient climatic cycles, the present-day warming takes place on decadal scales which needs to be interrupted immediately to maintain warming by maximal 1.5 °C (IPCC, 2022). Chapter 6. Sharp increase of lake temperatures over one decade during interglacial MIS 7c: Insights from the Dead Sea basin Thus, it is important to understand which short-term conditions (e.g., temperatures)

might have prevailed at terrestrial study sites in the past.

The Levant region, located in an arid climatic zone between the south-eastern edge of Europe, western Asia and north-eastern Africa, is an interesting area of study due to its climate sensitivity and archaeological Ąndings indicating an early inĆuence of our ancestors on the region in the last 300 ka (Slon et al., 2014). It was also probably the main corridor for migration to Asia and Europe (Hershkovitz et al., 2018).

The DSDDP core 5017-1-A retrieved from the Dead Sea basin (DSB) yields a full sedimentary record of the past 250 ka, thus, including several glacial-interglacial periods that are represented by speciĄc lithological units during wetter (detrital material) and drier (autochthonous gypsum, aragonite, salt) regional climate conditions. Halite units accumulated during lake mixing periods (monomixis) and respond to seasonal shifts of summer-winter in the present-day Dead Sea, which makes it attractive for short-term climate studies. Furthermore, previous studies have shown that Ćuid inclusions trapped in salt (halite, NaCl) serve as accurate palaeothermometer using Brillouin spectroscopy (Guillerm et al., 2020;Brall et al., 2022).

The work performed in this chapter deals with annually-deposited halite layers from a sediment core section that corresponds to the interglacial MIS 7c. Trapped Ćuid inclusions in coarse halite layers are used to determine their entrapment temperatures on decadal scales. The temperature difference of sub-samples in each layer will be used to calculate palaeoseasonal variations, and averaged annual temperatures will be compared to present-day temperatures as well as to the interglacial MIS 5e period.

Chapter 6. Sharp increase of lake temperatures over one decade during interglacial MIS 7c: Insights from the Dead Sea basin

Material

In this work, the oldest halite-bearing core section has been studied of core 5017-1-A. Core 5017-1-A has been drilled in December 2010 in the central basin of the Dead Sea (Levantine region, Israel) during the ŠDead Sea Deep Drilling ProjectŠ (DSDDP; Stein et al., 2011). The overall drill length was 1087.2 m, the cored length 922.4 m of which 78.2% could be recovered (that is, a core length of 721.2 m). For a general overview about the geological setting in the Levantine region including the DSB, the reader is referred to Chapter 5. Basically, core 5017-1 comprises two lithological groups (marl and halite facies, respectively) with main sub-facies of gypsum detritus (gd), alternating aragonite and silty detritus (aad), laminated detritus (ld), layered halite (lh) often accompanied with homogeneous halite (hh), and halite and detrital marl (hdm). Further, slumped deposits of mass transported deposits (mtd), halite transported deposits (htd), and coarse clastic detritus (ccd) have been deĄned (Neugebauer et al., 2014;Figure 6.2 A).

Core section 5017-1-A-176-2 belongs to the Amora formation and has been deposited

during Marine Isotope Stage 7 (MIS 7), located between 445.03 meters below lake

Ćoor (m blf) to 446.50 m blf (Figure 6.2 A). The lithology of section 176-2 is dominated by detrital dark clay layers (144.5 to 82 cm, 78.5 to 67 cm, and 6 to 0 cm of core section depth), an ochre-colored aragonite part of 3.5 cm total thickness (82 to 78.5 cm), and salt alternations (67 to 6 cm) (Figure 6.3 A, B). Thus, based on a core section length of 144.5 cm, the total fraction is largest for clay (55.4%), followed by salt (42.2%), and aragonite (2.4%). The closest dated tie point of 176-2 was published at 455.89 m blf with an age of 220 ± 2.0 ka (Goldstein et al., 2020). Another tie point, of the same study, was deĄned at 419.83 m blf with 207 ± 2.0 ka. Thus, both tie points lie 10.06 m below and 27.34 m above the lowest salt layer, respectively (Figure 6.2 B).

clay intervals are considered to represent meromictic lake regimes (no mixing throughout a year) due to modern observations in the Dead Sea (Gertman and Hecht, 2002). On the contrary, salt crystals have been reported to nucleate/deposit uniquely during holomictic periods when the water body mixes at least once a year (Gertman and Hecht, 2002). degree of supersaturation with respect to NaCl (Sirota et al., 2017). CO crystals nucleate from the brine below the thermocline in the deeper lake environment (hypolimnion) during the summer season, while dissolution takes place at the shallow lake Ćoor (epilimnion) (Sirota et al., 2017). During winter when the water body is mixed, cumulate crystals precipate uniformly in the epi-and hypolimnion by settling on the previously formed CO layers.

Core section 176-2 encloses 13 coarse and 12 cumulate layers of such halite types (Figure 6.3 A). Above CO layers the grain size boundaries show frequently sharp changes in both color and grain size, pointing to a rapid accommodation of settled CU crystals that form smooth surfaces (e.g., CO9-TOP -> CU in Figure 6.3 B-3; Kiro et al., 2016;Sirota et al., 2017). The transition of CU to CO types show in most cases also a sharp boundary (e.g., CU -> CO10-BOT in Figure 6.3 B-3). However, the exact position where CO crystals started to nucleate on the underlying CU layer is sometimes not easy to detect due to transitions zones of unconsolidated CU surfaces (e.g., CU -> CO2-BOT in Figure 6. The reader is referred to Chapter 5.3 for detailed information about the equipment used and the basic principles of Brillouin spectroscopy. Data are presented here as mean T x values for each sub-section (i.e., uppermost samples in coarse layers labelled "CO-TOP", and lowermost samples in coarse layers labelled "CO-BOT", respectively). The depth of the water column at time of halite deposition must be taken into account because the pressure affects the water density at the lake Ćoor (Guillerm, 2019). Due to the lack of a lake-level model, no correction term could be applied on measured T x for this study so far. Therefore, the presented results

show only the evolution of the relative lake bottom temperatures within 13 years of coarse halite deposition. An additional modelling study is in work to reconstruct the prevailing lake-level during MIS 7 by using the obtained solid-liquid-vaporequilibrium (SLVE) sound velocity values (Guillerm et al., work in progress). Still, we note that, in view of the reduced time interval, the lake level and composition was virtually constant during deposition of the studied samples. Therefore, the correction is the same for all data, and differences between values can be directly discussed without knowing the correction.

Chapter 6. Sharp increase of lake temperatures over one decade during interglacial MIS 7c: Insights from the Dead Sea basin top-related sub-samples reveals another interesting feature of core section 176-2.

In I low the CO-TOP sub-samples Ćuctuate constantly with local maxima (Lmax) during three-year cycles, including the transition to I cen from halite layers 13 to 8 (Figure 6.4). However, CO-BOT sub-samples are characterised by local minima (Lmin) every three years throughout the studied core section, with CO-BOT being in anti-phase to CO-TOP along the layers CO13 to CO7. Although the anti-phase is just showing the (weak) seasonal signal that had been deĄned earlier, the CO-BOT oscillation pattern changes from anti-phase to a phase difference during I cen to I upp (Figure 6.5).

This "parallel-like" oscillation along the layers 7 to 1 is described by: Chapter 6. Sharp increase of lake temperatures over one decade during interglacial MIS 7c: Insights from the Dead Sea basin

• an increase in both, CO-TOP and CO-BOT T x values (CO7-TOP Lmax; CO6-BOT Lmax),

• a two-year decline in both, CO-TOP and CO-BOT T x (CO5-TOP Lmin; CO4-BOT Lmin),

• by a successive rise in T x within three years (CO2-TOP Lmax) and two years (CO2-BOT Lmax),

• and by a Ąnal decline in T x in both CO-BOT and CO-TOP.

Total maximum T x differences (absolute maximum minus absolute minimum of each sub-sample type) are 3.4 °C (CO-TOP) and 5.2 °C (CO-BOT). The temporal differences (uppermost T x minus lowermost T x for each sub-sample type) are 2.3 ± 0.6 °C (CO-TOP) and 3.9 ± 0.6 °C (CO-BOT), respectively. Finally, mean temperatures increased by 5.0 °C (CO1-TOP minus CO13-BOT) during more than one decade of halite deposition, with a combined uncertainty of ± 0.9 °C by taking into account the uncertainty (SE) of both sub-samples (± 0.5 °C for CO1-TOP, and ± 0.7 °C for CO13-BOT, respectively).

The major results observed are:

• in three layers, the coarse halite derived (bottom) temperature is nearly equal or larger than top temperature

• in the Ąrst half of the layers, the top temperature is initially rather constant, and shows an increasing trend by around 3 °C in the second half

• the bottom layer oscillates throughout the layers, but shows an overall increasing trend by around 3 °C too

• similar temperature trends (rising/declining) of both top and bottom coarse halite samples from layers 8 to 1 Chapter 6. Sharp increase of lake temperatures over one decade during interglacial MIS 7c: Insights from the Dead Sea basin 6.5 Discussion 6.5.1 Meromictic vs. holomictic periods in the DSB clay deposits in the Dead Sea sedimentary record have been interpreted to reĆect rather wetter climate conditions associated with colder temperatures and increased lake levels (Torfstein et al., 2015;Kiro et al., 2016). A possible cause might have been increased moisture in the DSB due to diverted southward storm tracks with more intense and frequent Eastern Mediterranean cyclones (Torfstein et al., 2015). The overlying clay layer in core section 176-2 indicates thus a temporal shift towards wetter climate conditions in the Levant, while elevated temperatures in the studied uppermost halite interval reĆect the peak of a temporal dry period, which is in line with general interpretations of halite units in the Dead Sea basin (Neugebauer et al., 2014;Kiro et al., 2016).

The record of 176-2 provides a seasonal signature after the onset of halite deposition, with a calculated mean value of 2.7 °C (Figure 6. i.e. at the same frequency or amplitude of occurring natural hazards, because the regional climate has often a major control on environmental processes (IPCC, 2022). However, it has already been reported that the Northern HemisphereŠs seasons have shifted since 1952, and that summer days became longer and warmer while winters shortened but warmed too (Wang et al., 2021). Therefore, the interglacial period MIS 7c may help us to better understand on which time scales climate-sensitive continental regions underwent signiĄcant temperature changes or not. To identify potential seasonal changes over a decade and expecting a cooling (warming) during winter (summer), the temperature changes were calculated from data of both summer-to-winter and winter-to-summer intervals (see Figure 6.10):

∆T cooling = T x CO-T OP (n) -T x CO-BOT (n+1) ∆T warming = T x CO-T OP (n) -T x CO-BOT (n) (6.1)
Hence, typical winter seasons in the Levant would yield ∆T cooling values greater than 0 due to a drop in temperature (Gertman et al., 2010). Anomalies would be represented by ∆T cooling values less than 0 due to an increase in temperature during cumulate halite nucleation. Temperatures usually increase during spring to late summer when coarse halite crystal growth is dominating with ∆T warming values greater than 0. ∆T warming values less than 0 would rather imply a temperature anomaly.

Two cooling anomalies (∆T cooling greater than 1 °C) were identiĄed in layers 9/8 and 3/2 which is supported by one warming anomaly along the growth direction of the corresponding coarse summer layer 8 and ∆T warming close to 0 in layer 2 (Figure 6.7 D). To summarise, highest ∆T warming values appear on a 3-year cyclicity over one decade with a declining trend starting at the central part of the core section (layer 7).

The results imply that a similar seasonal shift like today could have occurred in MIS 7c when lake bottom waters became warmer in both, summers and winters. It is supported by slightly increased cumulate layer thicknesses (Figure 6.7 B), indicating higher degrees of supersaturation of salt like in the modern Dead Sea (Sirota et al., 2017). The timing of present-day lake mixing depends on the salt concentration in the upper water layer (epilimnion): at some point, the water density is high enough to induce Šdouble diffusion diapycnal ĆuxesŠ (DD Ćux) where salt is transferred to the deeper water layer (hypolimnion) (Arnon et al., 2016). Hence, the DD Ćux could have appeared earlier than usual (late November) in MIS 7c to account for a higher amount of nucleated cumulate crystals which form thicker layers at the lake Ćoor. However, the results obtained here do not allow to claim if the summers became even longer and the winters shorter, as observed for the Northern Hemisphere during the last decades due to modern global warming (Wang et al., 2021). The total mean temperature difference between the oldest and youngest coarse halite layers is 5 °C ± 0.9 °C. However, since a strong seasonality lasting for multiple years has been established ca 6 years after the onset of the holomictic lake regime, it remains questionable if a comparison is reasonable between layers CO1-TOP and CO13-BOT because the seasonal temperature difference might likely be included. Therefore, it is recommended to compare only data that are related to summer or winter sub-samples. Chapter 6. Sharp increase of lake temperatures over one decade during interglacial MIS 7c: Insights from the Dead Sea basin 6.5.3 Temperature shift for entrapment pressure

The analytical approach of measuring directly the entrapment temperatures of Ćuid inclusions using Brillouin spectroscopy is valid for (sub)surface minerals or rocks which formed close to ambient pressure. If this applies, then we write Tx = Tf. However, pristine evaporites can be found in even deeper environments in both oceans and lakes at higher water pressure. This pressure acts on the Ćuid inclusions during crystal growth, leading to a Ąnal temperature shift of both, "true" and measured temperatures. Hence, it is necessary to estimate either the water pressure at time of Ćuid entrapment (crystal growth), the water level (i.e., sea or lake level), or the water height (water level minus Ćoor level). It follows:

C = Q * D * g * H (6.2)
with C the temperature correction (°C), Q the quasi-isochore (°C/MPa), D the brine density (kg/m 3 ), g the gravitation acceleration at Dead Sea latitude (m/s2), and H the water height (m). For core section 176-2, the temperature correction equals (calculated by Guillerm et al., 2022, in

C (°C) = 0.74 °C/MPa * 1218.9 kg/m 3 * 9.795 m/s 2 * 318 ± 40 m C (°C) = 2.9 ± 0.4 °C Note that C is around 1.9 °C lower during MIS 7c than C during MIS 5e (chapter 5), despite similar lake levels (-306 ± 23 m vs. -311 ± 18 m, respectively). As an important factor to estimate the absolute water height, the subsidence rate of the DSB needs to be considered, and thus the Ćoor elevation. The subsidence rate during the last 220 ka has been estimated to 2.00-2.45 mm/year with minor variations (Goldstein et al., 2020). Hence, the subsidence rate might have likely outpaced the sedimentation rate of the basin, leading to deeper basin Ćoors during MIS 5e than MIS 7c. Therefore, both the water height and the pressure were higher as well, causing a higher temperature correction of 4.8 °C for MIS 5e samples (Guillerm, 2019;Guillerm et al., 2022, in review).

Final corrected Ćuid inclusions temperatures (in °C) using BS are calculated following:

T f = T x + C (6.3)
on each T x value from MIS 7c, and thus provides a straight comparison to other measured lake-bottom temperatures in the Dead Sea basin (Figure 6.8).

Chapter 6. Sharp increase of lake temperatures over one decade during interglacial MIS 7c: Insights from the Dead Sea basin

Palaeoclimate implications of Ćuid inclusion temperatures during MIS 7c

The applied pressure-related temperature shift on each T x data point provides a better comparison with both modern and MIS 5e data of the Dead Sea basin.

The resulting T f data are in relatively good agreement (50 % of the data within the error bands) with the averaged temperatures obtained from core 122-2 dating back to MIS 5e (gray band in Figure 6.8; Table 6.2). The lowest temperature in sample CO13-BOT (17.7 °C at 445.695 m blf) is indicative for a change of the lakeŠs hydroclimate which is associated with a lithological change from clay to salt deposition (Figure 6.3). This provides Ąrst evidence that such clay-dominated core sections may be associated with colder, wetter periods due to an increased runoff of Ąne detrital material into the lake, as it was interpreted in Ąrst studies of the environmental implications of the DSDDP sediment core 5017-1 (Kiro et al., 2016).

However, temperature estimates of such interglacial periods were missing until now due to the lack of accurate palaeothermometry methods to be applied on DSDDP The blue band displays the mean of corrected Ćuid inclusion data from core 5017-1-A-122-2 (MIS 5e) (error margin of ± 0.5 °C), the purple band marks the mean of core section 176-2. Note that the error overlaps with the gray band due to almost identical values. This Ąnding is in good agreement with previous interpretations of salt facies in DSDDP core 5017-1: salt facies were interpreted as regional responses to periods of global climate warming with drier, hypersaline conditions in the Levant region, associated with major lake level drawdowns (Kiro et al., 2017;Torfstein et al., 2015).

These arguments could also be used here to explain the maximum temperatures at the end of halite deposition in core section 176-2. However, recent results from analyzed Ćuid inclusion data (including the data set presented in this chapter), combined with reconstructed lake levels, subsidence rates, and the modeled NaCl inventory of the DSB (Guillerm, pers. comm.), suggest that lake levels were not primarily affected by global climate changes. The extent to which lake levels are correlated with (hypolimnion) temperatures remains an open question, which requires future studies using multiproxy approaches.

Nevertheless, the results suggest the great potential of analyzing halite section in the sediment record to estimate the thermometric evolution of the DS water body through time. Further, the method is a powerful tool to accomplish high-resolutional reconstructions of decadal palaeoseasonal variations in halite sections. In this study, Chapter 6. Sharp increase of lake temperatures over one decade during interglacial MIS 7c: Insights from the Dead Sea basin the estimated temperature increase over 13 years is about 2.3 °C, being a reasonable value compared to modern global warming in many continental regions (IPCC, 2022). This study suggests furthermore that continental water bodies such as the Dead Sea basin might experience decadal-to-centennial temperature Ćuctuations as seen by the relation between both, the onset (lowest T f ) and the end (highest T f ) of halite deposition during relatively drier interglacial periods. Lake bottom temperatures are estimated to have increased by 3.9 ± 1.4 °C for winter intervals, and 2.3 ± 1.2 °C for summer intervals, respectively. This is in contrast to present-day model projections that have shown that the Eastern Mediterranean region (including the Levant) will be affected by a strong warming:

Mean surface temperatures in winter seasons will increase by ca 2-3 °C between 2040-2070, and in summer seasons by ca 4-5 °C (Lelieveld et al., 2012)(Figure 6.9).

Another model for the Levant region reconĄrmed that global warming will result in a stronger warming of summers than of winters due to the westward expansion of a thermal low, which will increase the number of hot days per year, decrease the regional rainfall amount and lead to dry soils with less heat storage capacities (Lelieveld et al., 2016). The contradicting reversed decadal rise of both winter and summer lake temperatures could be explained by the onset of a global interglacial period (ca 220-210 ka; Martinson et al., 1987). Summer insolation was highest at that time during the past 400 ka [START_REF] Berger | Long-Term Variations of Daily Insolation and Quaternary Climatic Changes[END_REF], further evidence of a short-term (ca 16 ka; Tzedakis et al., 2004) global warming event comes from other proxy records, e.g., from the Antarctic Vostok ice core (Petit et al., 1999). However, the MIS 7c interglacial was not as pronounced as other strong interglacials because the peaks of both global mean temperatures and CO 2 concentrations were lower than MIS 11 or MIS 5e, although both variables increased in a short time (Figure 6.1;Petit et al., 1999). Hence, the lower CO 2 concentration during MIS 7c could have controlled the continental temperature effect like it does today, and explain why winters warmed stronger than summers in the Dead Sea Basin. Desprat et al. (2006) reported that Atlantic sea-surface temperatures increased by ca 8 °C in winter and ca 10 °C in summer, respectively, during the transition from MIS 7d to 7c.

Chapter 6. Sharp increase of lake temperatures over one decade during interglacial MIS 7c: Insights from the Dead Sea basin

Conclusions

A detailed palaeoclimate study covers an episode of dry conditions in Lake Amora (Dead Sea basin) during MIS 7c. The data set obtained in this study provides for the Ąrst time direct temperature measurements of the DSB over more than one decade. The record bears seasonal varying signatures along a halite unit that is comprised of 13 alternating coarse/cumulate layers. The presence of such halite layers between two thick clay units implies holomictic lake conditions at the time of deposition. Contrary to the onset of holomictic conditions in 1996 and a seasonality of 1 °C, the results presented here suggest a rather sluggish onset of annual seasonality (winter-summer) during MIS 7c. However, winter temperatures vary with a 2-3 year periodicity and rise steadily throughout more than one decade.

A general decadal temperature rise of lake-bottom temperatures can be inferred from the data set, as seen by averaged increases of both, 3.9 ± 1.4 °C for winter intervals, and 2.3 ± 1.2 °C for summer intervals, respectively. The overall mean temperature increase can be estimated to 5 ± 0.9 °C. The rapid warming of both winter and summer seasons imply that the Dead Sea Basin has ever been a very sensitive hydrological system, with a strong response of the dominating lithological facies to intervals of global climate warming. The sharp increase of palaeo-lake temperatures also suggest that the present-day global warming might strongly effect the yet hypersaline Dead Sea and adjacent water resources, which could exacerbate the regional water supply for people, and increase the vulnerable situation of the Levant.

This study conĄrmed the power of halite layers as climate archives when combined with Brillouin spectroscopy to determine palaeotemperatures. Future studies could be thus performed on other halite units in DSDDP core 5017-1-A to investigate palaeoclimate lake conditions for any dry period of interest.

Chapter 6. Sharp increase of lake temperatures over one decade during interglacial MIS 7c: Insights from the Dead Sea basin 6.7 Appendices

DeĄnition of varves

Formation temperatures of coarse halite samples have been used to compute the seasonal difference of individual layers by the offset between top T x data and bottom T x data (∆ T warming ) (see Chapter 5). It has been assumed that modern halite precipitation patterns might be valid for sedimentary records because the maximum palaeo-lake water temperatures were obtained for sub-samples from uppermost coarse intervals (CO-TOP). The modern Dead Sea is dominated by a holomictic lake regime with a mixing of the whole water body once a year (late November) due to salt Ćuxes from the lakeŠs surface to the lake Ćoor (Gertman and Hecht, 2002;Arnon et al., 2016). This process coincides with a change in the prevailing halite type from coarse to Ąne halite. Further, the depth in the water body, where halite nucleation takes place, turns from the lake Ćoor (coarse halite) into the whole water column (including the brine-air interface). Although other suggestions were made to deĄne lacustrine palaeoseasonality (Kwiecien et al., 2022), the approach using ∆ T warming may be most suitable in the case of annually-deposited halite layers in the DSB.

Two approaches to deĄne evaporitic varves have been used in this work to test the correlation between varve thickness and palaeotemperature data (Figure 6.10).

Initially, the thickness of individual halite layers has been determined for both types from the bottom to the top of the core section, by measuring the distance between the Ąrst clear occurrence of a halite type to the occurrence of the overlying other type (Figure 6.10 B). The documented position of each boundary has been used further to determine absolute varve thicknesses.

Couplets of two layers are termed Šwarming varveŠ by spanning from the coarse bottom to the cumulate top part of each layer (Figure 6.10 A). Hence, they include the period of initialing warming (CO-BOT -> CO-TOP during spring-summer) followed by a potential lake cooling (CO-TOP -> cumulate during fall-winter).

The second approach named Šcooling varveŠ uses couplets that are equal to initial lake cooling (cumulate during fall-winter) followed by a potential warming of the water body during spring-summer. Relative varves were calculated directly as differences (∆ V i ) between the documented location during sample preparation of each sub-sample type (Figure 6.10 A):

• CO-TOP(N+1) -CO-TOP(N)

• CO-BOT(N+1) -CO-BOT(N).

and are plotted against T x data in Figure 6.11. Note that the calculations of ∆ V i were carried out independently of the initial sampling at a later step of this work.

DeĄnition of temperature couplets

The large data set obtained from 13 halite layers with 26 sub-samples gives the unique opportunity to test any existing correlations between the relative ŠdistancesŠ of adjacent sub-samples and their proper ŠpalaeotemperaturesŠ. For that, T x data of all studied sub-samples were used to determine certain numerical differences ∆T i (Figure 6.10 C):

• ∆T annual pre-summer

• ∆T annual post-summer

• ∆T cooling

• ∆T warming Chapter 6. Sharp increase of lake temperatures over one decade during interglacial MIS 7c: Insights from the Dead Sea basin 6.7.3 Results

Absolute varve thicknesses

Data of all measured varves along the core section are listed in Table 6.7. Cooling varves show slightly varying values within I low from CO13-TOP to CO8-TOP between 4.0 cm and 7.2 cm with a mean thickness of 5.32 cm. After a local maximum of 7.2 cm (CO8-TOP), the data show a drop to 5.4 cm, 2.6 cm, and 3.4 cm, respectively, with a mean of 3.8 cm from CO7-TOP to CO5-TOP. The upper part yields values of 2.9 cm and 3.8 cm (CO4-TOP and CO3-TOP, respectively).

The last value obtained from CO2-TOP and CO1-TOP is the maximum value (7.7 cm) of all measured distances. The plotted curve against cumulate layer thicknesses correlates weakly due to the maximum at 7.2 cm (Figure 6.7 C).

Warming varves show only small variations compared to cooling varves (Figure 6.7 A). The lowermost interval of the core section yields three constant thicknesses (varves 11 to 9). An almost perfect correlation occurs during varves 8 to 6, subsequently the warming varves are in anti-phase to the cooling varves (varves 5-3), while the uppermost part of the core section is characterised by a parallel increase (Figure 6.7 A). Furthermore, the warming varve thickness correlates well (R 2 = 0.92) with the thickness of cumulate layers (Figure 6.7 B, C). The calculated relative thicknesses between CO-BOT sub-sample locations (∆V bot )

show a different evolution along core section 176-2. ∆V bot varies stronger in its amplitude with alternating drops and rises on two-year cycles from CO13-BOT until CO8-BOT with 3.0 cm < ∆V bot < 6.5 cm (Figure 6.11 A). A clear increase in ∆V bot with a total maximum (8.0 cm) along the core section occurs towards CO7-BOT, followed by an overall sharp drop with mainly declining ∆V bot values of 5.5 cm, 3.0 cm, and 2.5 cm until CO3-BOT. The two uppermost sub-intervals yield slightly raised thicknesses of 4.3 cm and 3.9 cm. The main Ąnding here is that both obtained curves of ∆V bot and CO-BOT T x are Ąrst in anti-phase until Chapter 6. Sharp increase of lake temperatures over one decade during interglacial MIS 7c: Insights from the Dead Sea basin halite layer CO7, and secondly in phase until layer CO1 (Figure 6.11 B). [cm]

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Introduction

The PhD thesis was initially planned to conduct experiments using Brillouin spectroscopy on lacustrine gypsum samples from Lake Petén Itzá (Guatemala)

and the Aral Sea (Kazakhstan, Uzbekistan; see chapter 2.1). Both locations are sensitive to climate variations on shorter time scales, which can be seen in their sediment records with a response of the dominating mineralogies to colder/drier (gypsum, salts) or warmer/wetter (carbonates, clay) climates (e.g., Sorrel et al., 2006;Oberhänsli et al., 2011;D. Hodell et al., 2006;D. A. Hodell et al., 2008;Hillesheim et al., 2005). This rock-climate relation offers a great potential to study abundant gypsum intervals.

The experimental part of applying BS on such samples has started with initial essays on modern gypsum crystals from a salt works in Trapani (Sicily; provided by M. Natalicchio). Such material proved to form in equilibrium with the saline parent brine during gypsum growth with the entrapment of Ćuid inclusions in growth bands along the host crystal direction of growth (Natalicchio et al., 2014. Although such growth bands with triangular shapes with Ćat bases of FIs were observed, the material did not allow to perform BS measurements, because:

• the FIs were mostly destroyed during sampling preparation (not liquid)

• the few liquid FIs analyzed did not provide clear Brillouin spectra

• the cavitation into biphasic FIs was not successful in most cases Therefore, the material was discarded to set a "gypsum standard" for BS, and laboratory-grown gypsum has been further used throughout the experimental work during the PhD thesis.

Furthermore, it has been noted, after more than one year of experimental work during this thesis, that such lacustrine gypsum samples are not suitable for the BS approach, for several reasons:

• crystals often have lenticular shapes (difficult to prepare samples)

• crystal sizes are usually less than 500 µm

• low abundance of trapped FIs

• gypsum transforms rapidly into hemi-hydrate phase upon warming

• low amount of cavitated FIs (necessary for T x measurements)

• <10 T x points were measured in two years, ranging between 0-60 °C

Cavitation of gypsum Ćuid inclusions

The presence of a vapor in a liquid inclusion is the essential second step during experimental BS analysis, as in this case for gypsum. In general, a heating cycle is preferred to cavitate a vapor because it works without the use of liquid nitrogen, and is thus more economic. For gypsum, however, the previously proposed shortterm heating to 130 °C for halite samples (Guillerm, 2019) can not be applied due to a fast thermal response of the mineral to elevated temperatures (T ) which is associated with a loss of the molecular water, and a transformation to the hemi-hydrate phase bassanite (Figure 7.1). Since each T x value depends on the slope of both isochores and LVE curves,

Natural gypsum

A total of 32 T x data could be obtained by using natural gypsum samples for BS. Unlike for the challenging experimental work using lab-grown gypsum, Ćuid inclusions trapped in natural samples were generally easier to use because of higher probabilities in cavitating vapor bubbles during cooling-heating cycles (see section thermal cycles). Furthermore, a sufficient amount of FIs inside any sub-sample has always been found. In some cases, secondary FI types were used for T x measurements (e.g., samples from Santa Elisabetta; Eraclea Minoa; Aral Sea). For such cases, Ątted isochores / LVE points are mainly crossing with gently sloped LVE curves, resulting to reasonable T x values (Figure 7.5).

Discussion

Studies dealing with the determination of formation temperatures of natural/labgrown halite have already shown that Brillouin spectroscopy is a suitable technique with an uncertainty of ca 1.5 °C (see chapter 5;Guillerm et al., 2020). However, no studies have been performed on gypsum yet, which is after CaCO 3 the second mineral that precipitates from an evaporating brine or an ideal seawater solution (Holser, 1979). Hence, gypsum forms usually in surface brines supersaturated with respect to calcium sulfate and close to ambient pressure, which makes this mineral favorable for palaeotemperature-related studies.

However, there are some properties of gypsum formed in natural environments that are limiting its usage for such studies, because

• gypsum has a low temperature-stability Ąeld (<42 °C) that enables dissolutionreprecipitation with time

• a MohsŠ hardness of 2 can trigger the deformation of both the host crystal and the trapped Ćuid inclusions,

• the high crystal water content (up to 20%) makes it favorable for dehydrationrehydration processes

• Ćuid inclusions can derive from several post-entrapment generations,

• former primary FIs trapped along the growth bands are vulnerable for solidiĄcation with time,

• FIs in gypsum often fail to cavitate a vapor or need multiple cooling-heating cycles.

Furthermore, the low solubility of CaSO The most important negative point regarding the lab work is the low abundance of cavitated gypsum FIs after the measurement of isochoric data point had been done.

To conclude one can state that cavitation occurs more often in natural gypsum than in lab-grown gypsum, and more often in secondary than in primary FIs.

Besides the above mentioned points, an additional limiting aspect by using gypsum

FIs for Brillouin analysis is that the acquired ∆f B spectra often can not be Ątted despite a sufficient amount of counts (≥300), if the inner Ćanks of the ∆f B spectra are too narrow to the Brillouin reference peak. In this case, the analysis program can not read out the ∆f B peak, causing high chi 2 values with w < 1000 m/s or >2500 m/s. One solution might be to deconvolute the FI spectra or locate its peak value, which could require a manual adjustment for each sample. Ćat SLVE curves as it has been observed for the accurate BS application on natural halite samples from the Dead Sea (Guillerm, 2019). It is likely that crystal thicknesses play a role in acquiring well-Ątted curves of FIs because the latter is directly related to the depth of FIs relative to the crystal surface.

A last point to be improved in the future when using BS on gypsum samples would be establish a gypsum-reference sample to correct for the instrumental drift during an experimental day in the laboratory. This is simply due to the fact that no suitable gypsum-reference crystal could be produced so far which would be big enough to work with on a daily basis. This aspect should be taken into account when applying BS on gypsum samples, because the usage of a NaCl-reference sample can cause problems during data analysis due to an increased frequency shift in saturated NaCl (ca 9 GHz vs. 6-7 GHz in low-saline gypsum). It is further recommended to use large-sized selenite samples from modern salt works, combined with in situ measurement of the physical-chemical properties of the parent brine.

Summary

For the Ąrst time, the calcium sulfate mineral gypsum has been studied using Brillouin spectroscopy (BS) to determine entrapment temperatures of Ćuid inclusions.

The experiments revealed a challenging data acquisition in lab-grown material which is needed initially to test a new methods to known standard conditions.

The general amount of trapped Ćuid inclusions in gypsum crystals was low which prohibited a complete study to setup the basis for accurate measurements of geological records.

Nevertheless, material has been used from natural subsurface locations, and trapped Ćuid inclusions provided a slightly larger data set to get Ąrst insights into BS analysis on this evaporitic mineral. However, the data are mainly derived from secondary FIs, which questions if the mineral is suitable to determine the pristine formation temperatures of gypsum samples. Given the potential danger of post-depositional sampling alterations in gypsum and its relatively low MohŠs hardness of 2, one might however argue to avoid using gypsum for studies that aim to reconstruct palaeoenvironmental conditions at the time of crystal formation.

Nevertheless, the results rather suggest that BS could be used as a tool to study secondary FIs (e.g., the evolution of sealed cracks, the distribution of temperatures along FIAs).

More research is needed to fully answer this question, and future studies dealing with this matter should use largely grown samples obtained from locations of known growth temperatures. Further, the origin of trapped Ćuid inclusions should be determined to rule out post-entrapment alterations. A general recommendation is to focus on triangular-shaped FIs aligned on growth bands within a crystal which indicates pristine signals.

Introduction

It has been known for several decades that the speed of sound in aqueous solutions depends on variations in salt concentration, temperature and pressure (Chen et al., 1978;[START_REF] Kleis | Dependence of speed of sound on salinity and temperature in concentrated nacl solutions[END_REF]. Therefore, Brillouin spectroscopy might be a potential tool to reveal relative variations in palaeosalinity from Ćuid inclusions along the direction of growth and/or deposition in geological samples such as evaporites. For this reason, the scientiĄc goal of this chapter was to conduct Ąrst experiments on saline aqueous solutions and to compare the measured sound velocities with the calculated ones, and second to apply this approach on a natural gypsum sample from the sabkha El Melah, Tunisia. The selenite-type sample was selected due to its high abundance of growth bands with primary Ćuid inclusions along 6 cm, opening the opportunity to test if sound velocities vary signiĄcantly from the youngest to the oldest part of the host crystal.

The experimental part (saline aqueous solutions) aims also to deĄne a standard as a function of salinity from low-to high-saline water samples, while the practical part (gypsum sample) is intended to estimate the palaeosalinity by taking into account the Ąnal ice melting points of Ćuid inclusions too. In addition, both primary and secondary FIs were measured in the natural gypsum sample to compare several FIAs, and to detect differences in sound velocities that might be related to variations in salinity upon crystal growth. [1] and calculated [2] sound velocity data of pure H 2 O between 5-65 °C. [2] after Chen and Millero, 1986. The offset is shown as the difference of [2] and [1], correction factor as the ratio of [1] : [2] , and the error represents the deviation of [1] to [2].

T w [1] w [2] (Chen and Millero, 1986). Sample concentrations of dissolved NaCl indicated by salinity S and molality m.

S m w

[A] calc.

w [B] meas.

offset c Error after 4.6 capillaries The sound velocity depends on the salinity of any liquid often with linear relationship (Chen and Millero, 1977;Brewer et al., 2015). The derived BS data for low-saline ranges (0-5 wt.% NaCl) show a parallel pattern of increasing w as a function of T and S, which is in good agreement with modelled Ąts by Allen et al. (2017) sabkha El Melah (Tunisia), shoreline near to Ein Gedi, the Dead Sea, and two ponds with groundwater outĆow at the western shoreline of the Aral Sea ("AS11", "AS13"). The sample AS11 (or "Aral11") not only has the highest w values, but is also the only one of all measured samples to have a Ątted sound velocity curve that decreases strongly with increasing temperature (Figure 8.5). All other samples always increase towards a certain local maximum, followed by w values that rather form a plateau than a negative slope like Aral11 (Figure 8.5, ). sample "07" (-63 °C) while those of "05" and "09" are equal (-54.7 °C). The eutectic points are relatively low compared to that of NaCl-H2O (-21.3 °C), pointing to it being a system of more than two components, however, the signiĄcance of such partly erroneous observations should also be taken with a degree of caution (see discussion).

[A]-[B] [B]:[A] c [wt.% NaCl] [mol/kg] [m/s] [m/s] [m/s] ±0
Upon further slow heating, a vapor occurred in most cases around -45 to -40 °C, while FIs with small (<10 µm) sizes needed two or more cooling-heating cycles to nucleate a vapor. Lowest T ice values were observed in sub-sample "07" (-35.7 °C).

In general, the obtained overall data set covers a temperature range between -7.1 to -35.7 °C (Table 8.4). Both mean T ice and the maximum negative T ice values for each sub-sampled data set show a trend towards less negative values along the direction of growth of the host crystal (Figure 8.9, Table 8.4). The combined T ice vs. sound velocity plot indicates that both crystals "09" and "07" from the intermediate-oldest parts fall into distinct intervals (1600-1800 m/s and 1800-1900 m/s, respectively), as previously seen for (non-)saline capillary solutions (Figure 8.10 D). A relatively narrow T ice range is observed only for crystal "07" (Figure 8.10 D). On the contrary, T ice data of crystals "05" and "09" form a broader interval between -7.1 to -32.4 °C, and -12.6 to -35.3 °C, respectively. The enlarged interval might be likely due to internal growth band distances from up to 1 mm. Any observed data outliers were checked using the microscopic images of each distinct biphasic FI phase to identify high ratios of vapor:liquid in small FIs. A vapor bubble Ąlling the entire FI space affects the Brillouin spectra due to the interference of the laser light with the vapor phase. A crucial point of determining the bulk salinity of Ćuid inclusions is the identiĄcation of the correct chemical system (Goldstein and Reynolds, 1994).

A problem is that the true chemical composition of Ćuid inclusions (FIs) or a selected Ćuid inclusion assemblage (FIA) remains unknown at higher salinities without additional analytical methods. In the case of the Messinian Salinity A 445 m long sedimentary core from the DSB showed depositional features of varved halite units formed under increasingly warmer global climate conditions.

Such interglacial periods are coherent with less negative isotopic excursions on global scales (Marine Isotope Stages, MIS). Modern halite deposits respond to seasonal changes by precipitating Ąne cumulate (winter) or coarse (summer) halite crystals in the whole water column (Ąne) or solely on the deep lake Ćoor (coarse).

Thus, the modern Dead Sea served as an analogue for interglacial periods with halite deposition.

In this work, a new approach to study varved halite section has been applied to decipher palaeo-seasonal temperatures during two interglacial periods (MIS 5e, MIS 7c). By preparing sub-samples at lowermost and uppermost parts within individual halite layers, differences in palaeotemperatures were studies along the direction of growth. Thus, palaeo-seasonal periods of a single year are reĆected by sub-sampling locations in coarse layers during late winter (lowermost parts) and late summer (uppermost parts).

The study shows a cyclic increase in palaeotemperatures of coarse halite along the growth direction in three consecutive varves (= three years) at the beginning of MIS 5e. In contrast, data from Ąne-grained cumulate halite layers reĆect consistently lower temperatures. The inferred summer-winter seasonality was about twice as large as today. However, the number of samples is too small to draw conclusions about long-term trends in Lake Samra. Therefore, a salt interval of another interglacial (MIS 7c) was selected to reconstruct for the Ąrst time the complete temperature development of the deep lake layers over a period of 13 years. From the lithology of the drill core, it can Ąrst be concluded that the salt layers surrounded by clay layers were deposited in holomictic periods. Data from coarse-grained layers show that summer temperatures initially remained unchanged for 5-6 years after the start of the annual mixing process, whereas winter temperatures varied. A stable annual seasonality with clear temperature differences of about 2 °C between summers (warm) and winters (colder) set in after about 6 years. Furthermore, a

Chapter 9. Final conclusions relatively high temperature increase of up to 5 ± 0.9 °C was observed over more than one decade. Overall, the mean winter temperature increase was 3.9 °C and Overall, the results obtained in this thesis proved that the approach to combine evaporites with the Brillouin spectroscopy method gives hope to investigate more well-dated geological records at varying temporal scales. This tool is not limited to other proxy methods, and therefore powerful palaeoclimatic reconstructions can be performed on salt records in any geographical area. Such high-resolution studies on halite intervals will be helpful to gain a better understanding on seasonal processes in lakes for an area of interest. Here, it is recommended to use preferable halite instead of gypsum because practical work done in this PhD project has shown that large data sets can be obtained within a short time.

use carefully singlewise X2, Y2, X1, Y1.

Note that the position of Z might change quickly, so that you have to reset the peak again to 0 GHz. Afterwards press "stabilization on" at the interferometer control.

Wait several seconds (or minutes) for the signal to be computed and stabilized until it reaches ca. 4x times the signal intensity of the initial baseline.

10. Further, check the shape of the reference laser signal: This is very important before starting the experiment because the laser could not stay in monomode but in multimode. This has a direct inĆuence on all measurements and might be responsible for a full laser signal shutdown while analysing samples. Hence, the reference laser signal has to be sharp and symmetric, without having any bulbs or plateaus. In case of laser multimode, stop stabilization ("stabilization off"), go back to align mode ("align"), close the shutter, turn the big black knob at the laser control to 0.01 W, turn the key to the left, and wait for 5-10 minutes. Restart the laser afterwards and check again the signal shape.

11. Loading the sample: when you use single crystals, make always sure that a glass plate is already covering the silver block in the Linkam stage! Otherwise the sample can fall into the hole, which will result in much additional work to clean the whole instrument. When you use a capillary already stucked with wax on a glass plate, use an air siphoning tool: Ąrst, remove the big whitish tube from the Linkam and replace it, then use 1 LNP an push the small hole at the air siphoning tool for aspiration, remove the Ąnger to make it fall.

12. Measuring a ∆f B signal in a liquid: after having set the laser spot on your sample, switch from "shutter" to "window mode" at the interferometer control (Ąrst!) and set carefully the pinhole at the entrance to the black Brillouin box to "300". Pause the camera. While still observing, turn off the microscope light

ICI 2 Lmin 2 MN 1

 221 thank my parents Sabine and Peter, Stefan, Gerda, Kerstin, and Luca, for their mental and Ąnancial support when things got tight; Max, Nils, Alek and Jan for their their constant encouragement and moral support. Love you guysConĄdence interval to estimate the range of values that likely include the unknown mean of a sample population; here always referred to 95% with Z = ±1.96; thus the margin of error is 1.96 times the standard error of the mean CO Coarse coarse halite; grain sizes typically >1000 µm CU Cumulate coarse halite; grain sizes typically <500 µm CaSO 4 *2H 2 O Chemical formula of the sulfate mineral gypsum ∆d abs Absolute thickness of coarse or cumulate halite layers (cm) ∆f B Frequency shift between incoming and scattered laser beam by using Brillouin spectroscopy ∆T cooling Difference in T x between CO-TOP and overlying CO-BOT samples, representing the seasonal transition from summer to winter ∆T P Constant accounting for hydrostatic pressure on Ćuid inclusions; applied on T x values when estimates are available for the palaeo water depth ∆T warming Difference in T x between CO-BOT and overlying CO-TOP samples, representing the seasonal transition from winter to summer ∆V i Thickness of relative varves along coarse-cumulate halite layers by using the documented core section depth (cm) with i = top or bot DS Dead Sea DSB Dead Sea basin DSDDP Dead Sea Deep Drilling Program FI(s) Fluid inclusion(s) xv FI ref Reference Ćuid inclusion, measured every two hours while using BS. It is used to correct for the daily instrumental drift FIA Fluid inclusion assemblage FIB Fluid inclusion band I low Lowermost interval in core section 176-2 from 445.695 to 445.423 m blf (CO13-BOT to CO8-TOP); based on constantly low T x values I cen Central interval in core section 176-2 from 445.37 to 445.255 m blf (CO7-BOT to CO4-BOT); based on the onset of stable seasonality I upp Uppermost interval in core section 176-2 from 445.09 to 445.24 m blf (CO1-TOP until CO4-TOP; based on maximum T x values IRF Instrumental resolution function; used as a constant parameter of the Brillouin spectroscope ka / kyr Kilo years; or 1,000 years Lmax Local temperature maxima along coarse halite layers in core section 176-Local temperature minima along coarse halite layers in core section 176-Molar mass (g/mol) Ma Million years; or 1,000,000 years MIS 5e Marine Isotope Stage 5e (ca 130,000 -116,000 years ago); last interglacial MIS 7c Marine Isotope Stage 7c (ca 220,000 -190,000 years ago); interglacial MSC Messinian Salinity Crisis m blf Sediment core depth (meters below lake Ćoor) Number of analysed Ćuid inclusions or (sub)-samples N.A. Not determined NaCl Chemical formula of the salt mineral halite S Salinity (weight % NaCl) SD Standard deviation calculated for a sample population SE Standard error of the mean calculated from the standard deviation LIST OF TABLES SLVE / LVE Solid-liquid-vapor-equilibrium; term to describe data derived from biphasic FIs T Temperature (°C) T e Eutectic temperature; Ąrst melting point of ice crystals in a FI (°C) T h Homogenisation temperature of FIs: disappearance of vapor bubble (°C) T x Initial crossing temperature of single FIs by isochore and LVE (°C) T x mean Average of analysed FIs which are trapped in a sub-sample (°C) T f Pressure-corrected entrapment temperature of FIs by adapting a constant on T x (°C) T ice Ice melting temperature; last point of melting of ice crystals in a FI (°C) w Sound velocity (m/s) Historical background of palaeoclimate research Our modern understanding of complex interrelationships in the atmospheric Earth system is fundamentally shaped by multidisciplinary methods in geosciences. From the Ąrst insights into the four principles of stratigraphy and the detection of crystallographic structures in minerals by Nicolaus Steno in 1669, it took more than three centuries to develop climate models of the past and future. The study of past climate conditions on EarthŠs surface is one of the major keys to understand modern and anticipate future global climate conditions, as illustrated by the essential sentence "the present is the key to the past". It is related to the Ąrst Ąndings of modern geology in the 18th and 19th century by James Hutton and 42 years later by Charles Lyell

Figure 1 . 1 A

 11 Figure 1.1 A) 3D Sketch showing the principal growth of halite. A centered nucleus serves as starting point, followed by planar and diagonal growth along the edges. B) Top view on halite growth zones with trapped Ćuid inclusions. Red arrow marks the highest growth rates along the edges. Not to scale.

Figure 1 . 2

 12 Figure 1.2 Calcium sulfate minerals, in monoclinic crystal system and molecular water (reduced water) in gypsum (bassanite), and in the orthorhombic system after dehydration (anhydrite II, III) (from Schmid et al., 2020).

Figure 1 . 3

 13 Figure 1.3 Common shapes of Ćuid inclusions in minerals of halite (A, B) and gypsum (C, D). Note that three-dimensional images of FIs as in A-B images are scarce. Top-view as C) with monoclinic negative crystal shapes are usual. FIs can even be trapped along growth bands, as shown in gypsum from Sicily (D; Sample provided by M. Natalicchio). Many thanks to W. Ben Aissa for providing the images in A-B.

Figure 1 . 4

 14 Figure 1.4 Relation between conductivity, temperature, and salinity in global ocean surfaces (from Tyler et al., 2017).

Figure 2 . 1

 21 Figure 2.1 Sampling locations, marked by white rectangles, of material used for the PhD thesis. 1 = Banengo outcrop, Piedmont basin, Italy (Miocene; Messinian); 2 = Diverse outcrops on Sicily, Italy (Miocene; Messinian); 3 = Sabkha shoreline, Sebkha El Melah, Tunisia (Holocene); 4 = Bore core from the Dead Sea basin, Israel (Holocene-Pleistocene); 5 = Aral Sea shoreline, Kazakhstan (Holocene).

Figure 2 . 2

 22 Figure 2.2 (A) Map showing the arid climate regions of North Africa and Western Asia. (B) Dead Sea location with northern and southern basins.

Chapter 2 .Figure 2 . 3

 223 Figure 2.3 Overview on the remains of the Aral Sea (Red rectangle). Caspian Sea to the West. (B-D) Aral Sea satellite images showing the excessive water loss from 1973 (B), 1988 (C), and 2020 (D). Sources: Google Earth.

Figure 2 . 5 )Figure 2 .

 252 Figure 2.4 A) Study site of Sabkha El Melah in Tunisia. B) Modern-day evaporitic pan with halite nucleation (white colours). Note the proximity to the Mediterranean Sea. C) Sampling sites of gypsum material MG1.

Figure 2 . 5

 25 Figure 2.5 Principal geological units of the sabkha El Melah, Tunisia. ModiĄed afterDavaud et al., 1996.

Figure 2 . 6 A

 26 Figure 2.6 A) Messinian evaporites in the Mediterranean. A-1) Map of proposed sub-basins. CB, Caltanissetta basin; SB, Sorbas basin; VdGB, Vena del Gesso basin (Slightly modiĄed from Manzi et al., 2012, published Ąrst in Rouchy and Caruso, 2006). B) Overview on locations with Messinian deposits which were relevant for this work. C) Piedmont Basin. D) Outcrops on Sicily in the CB. Red star marks the location of a seawater sample.

  Laboratory experimental work was required initially to test the applicability of Brillouin spectroscopy on natural gypsum minerals from geological records. Thus, laboratory-grown (Šlab-grownŠ) gypsum crystals were produced from aqueous solutions under varying conditions (temperature, concentrations, degree of saturation, salinity, and wind velocities) to determine the most effective crystallization pathway for bulk crystals, by means of maximum crystal sizes with a maximum amount of entrapped FIs.All experiments implied a full evaporation of the aqueous solutions until the last solid crystals precipitated. The Ąrst evaporitic experiments were conducted by dissolving pure CaSO 4 • 2 H 2 O powder (Sigma Aldrich, ReagentPlus ® , ≥ 99%) in beakers (volumes = 50, 100, 500, 1000 ml) with pure H2O(Milli-Q ® , 18.2 M Ω cm). The solubility of CaSO 4 • 2 H 2 O in 1 L H 2 O is 2.2 gat 25 °C which is very low compared to NaCl (360 g/L). The Ąrst solution has been prepared to investigate the synthetical process of crystal shapes, sizes, and entrapment of Ćuid inclusion in solution of 1 L with 2.2 g dissolved CaSO 4 • 2 H 2 O powder. Gypsum nucleation started after evaporation of ca 220 mL H 2 O and continued until full evaporation as seen by Ąne layers on the beaker glass walls (Figure 3.1 A). After ca 10 days of evaporation at constantly 21 °C, very tiny (<20 µm) gypsum crystals with non-detectable thicknesses remained at the beaker bottom (Figure 3.1 B). The second solution ("Solution II") was prepared by dissolving 0.88 g CaSO 4 • 2 H 2 O in 0.4 L H 2 O, thus the solution was saturated by 96.1%. The powder was dissolved after 24 hours and the stock solution of 0.4 L further divided into smaller experimental solutions to test gypsum crystal growth at different temperature conditions. Solution II-A (volume = 0.1 L) evaporated for 1 week in a 1 L glass beaker using a thermal bath set to 26.5 °C. The actual beaker water temperature was constantly 25 ±0.4 °C. After complete evaporation of 0.1 L solution, the resulting gypsum crystals were Ąne-grained (<50 µm) with Ąbrous-needle shapes forming abundant circular assemblages. Solution II-B (volume = 0.187 L) was evaporated in a 0.2 L beaker contemporaneously on a stirring hotplate (Cat ScientiĄc MCS 66) with a temperature set to 60 °C. The beaker water temperature reached a constant value of 42 °C after 6 hours. First gypsum crystals were observed to appear at the water-air interface as small nucleates after 47% of H 2 O evaporation loss. The Ąnal solids were collected after 1 week.

Figure 3 . 1

 31 Figure 3.1 Lab-grown gypsum material of purely dissolved CaSO 4 • 2 H 2 O in water. A) nucleation steps of gypsum during evaporation of water, forming thin layers. B) Top view of (A) with tiny (<40 µm) crystals. C) Gypsum nucleation takes places at the air-solution interface (red arrow), and D) sinks then to the beakerŠs bottom (red arrow).

Figure 3 . 2 Figure 3 . 3

 3233 Figure 3.2 Microscopic view on synthetically-produced gypsum samples. A, C, D) precipitates from bulk solutions with dissolved CaSO4*2H2O in pure water. B, E) Crystals from bulk solutions containing small amounts of NaCl. Note the difference in crystal sizes between salty bulk solutions and pure water.

Figure 3 . 4

 34 Figure 3.4 Water samples used in this work for sound velocity measurements (11 = AS11, 13 = AS13). See chapter 8.2.3 for results.

  Figure 3.5) is the relatively youngest part, and the pyramidal-shaped part the relatively oldest part (lowermost part in Figure 3.5. Three sub-samples were cut out of the crystal along the { 010} cleavage plane by using a razor blade with labels 05 (youngest part), 09 (intermediate part), and 07 (oldest part). Each sub-sample contains three types of FIs: growth bands with aligned primary FIs having Ćat bases and triangular-like shapes to the top (Figure 3.6 A), randomly distributed elongated Ąelds with abundant occurrences of probably secondary FIs, and Ąelds of secondary FIs with smaller but broader irregular shapes (Figure 3.6 B).

Figure 3 . 5 Figure 3 . 6 Figure 3 . 7 Figure 3 . 8

 35363738 Figure 3.5 Selenite gypsum sample MG1 from the El Melah sabkha environment. Blue arrow indicates the direction of growth.

Figure 4 . 1

 41 Photon detector

Figure 4 . 2

 42 Figure 4.2 Brillouin spectra of a FI (red circle) in a lab-grown gypsum crystal. Left: Gypsum FI. Right: Recorded spectrum with reference line (violet) and inelastic (yellow) scattering from the FI.

4. 1

 1 photomultiplier, and to Ąt the resulting function on the raw spectra[START_REF] Pallares | Anomalies in bulk supercooled water at negative pressure[END_REF]. Three Ątting parameters are used in each analysis: c, τ , and an overall intensity factor K. The merit function χ 2 is minimised by the sum of square errors and normalised by the error bars on the data points. The error bars are caused by the photon noise which is deĄned as the square root of the number of counts[START_REF] Pallares | Anomalies in bulk supercooled water at negative pressure[END_REF].

Figure 4 . 3

 43 Figure 4.3 Fitted Brillouin spectra by using a speciĄc code written in Wolfram Mathematica.

Figure 4 . 4

 44 Figure 4.4 Left: Sketch of the cavitation process of a FI.A monophasic FI follows a quasi-isochore (dashed black line) upon heating (red arrow) with increasing internal pressure. During cooling (blue arrow) the internal pressure becomes negative. The FI stretches to to overcome the metastable phase (dashed black rectangles). Energy release forces a vapor to nucleate to re-enter the stable phase. By theory, the thermal path of a biphasic FI follows SLVE that represents ambient pressure conditions. Right: Sound velocity measurement in monophasic (black dashed line; isochore) and biphasic (red solid line; SLVE) phases by at four temperature steps. The intersection of both lines (T x ) gives the entrapment temperature (T f ) if the sample has precipitated at ambient pressure. In most geological cases, a pressure-correction is needed.

  7) with T ice the Ąnal ice melting temperature (in °C) of each FI. The resulting nonlinear Ąt is shown in Figure 4.5). Since the exact chemical compositions of the individual gypsum FIs could not be determined elsewhere, those FIs were Ąrst also analysed whose eutectic points were ̸ = -21.2 °C. .% NaCl]Final ice melting point[°C]

Figure 4 . 5

 45 Figure 4.5 Final ice melting temperatures as a function of salinity in salty aqueous solutions after Bodnar (1993).

  spectroscopy is used here to reconstruct annual resolution temperatures from a halite core section in the Dead Sea Basin during the onset of Marine Isotope Stage 5e (ca 130 ka) of the Last Interglacial. Lake bottom temperatures can be inferred based on the occurrence of coarse/Ąne halite facies, as observed today with the formation of equivalent halite facies during winter/summer seasons in the Dead Sea. A recurring increase in lake bottom temperatures is found along the direction of coarse halite layers in three successive years. Moreover, low Ćuid inclusion entrapment temperatures were detected in layers of Ąne (cumulate) halite facies.These results imply a two-fold stronger seasonality in the Dead Sea Basin compared to today, with colder winters at the onset of Marine Isotope Stage 5e. The results therefore highlight the potential of using cyclic salt deposits to reconstruct seasonal temperature variability for numerous evaporitic environments in the geological record. variability in the Dead Sea Basin during MIS 5e5.1 IntroductionPalaeoclimatic reconstructions often show that long-term global climate changes are largely induced through processes and feedbacks on seasonal timescales(Werner

  variations using continental archives to improve long-term palaeoclimate reconstructions.In this study, halite (NaCl) samples from the Dead Sea Basin (DSB) are used to reconstruct seasonal temperature variability during Marine Isotope Stage 5e (MIS 5e). Today, halite precipitates on the deep hypolimnion lake Ćoor of the DSB as alternating layers responding to a seasonal frequency(Sirota et al., 2021; Kiro et al., Chapter 5. Reconstructing lake bottom water temperatures and their seasonal variability in the Dead Sea Basin during MIS 5e 2016). During their formation, NaCl crystals may trap droplets of the surrounding lake water from which they precipitate, called Ćuid inclusions (FIs). Recently, a new palaeothermometer based on Brillouin spectroscopy (BS) was elaborated on FIs in natural quartz (Mekki-Azouzi et al., 2015) and laboratory-grown NaCl crystals(Guillerm et al., 2020), and used on a sedimentary core from the DSB to reconstruct temperature and lake-level changes of the Dead Sea over the Last Interglacial (LIG, 130Ű115 ka;Guillerm, 2019). In the present contribution, the focus is on a shorter time interval encompassing three closely spaced samples, to reconstruct palaeotemperature variability of the lake waters at an unprecedented (i.e., seasonal) resolution.Chapter 5. Reconstructing lake bottom water temperatures and their seasonal variability in the Dead Sea Basin during MIS 5e

Figure 5 . 1 (Figure 5 . 1 (

 5151 Figure 5.1 (Caption next page.)

  . The Ąrst long sedimentary core covering this time interval was retrieved within the framework of an International Continental Drilling Program (ICDP)-sponsored initiative (5017-1-A) (Neugebauer et al., 2014; Stein et al., 2011) (Figure 5.1 C). In this core, a layered halite facies (ŠlhŠ) accounts for around 20% of the sedimentary sequence and is composed of alternating grey marl Chapter 5. Reconstructing lake bottom water temperatures and their seasonal variability in the Dead Sea Basin during MIS 5e

Figure 5 . 2

 52 Figure 5.2 (A) Left, image of core section 122-2 with alternations of mud (M), cumulate (CU) and coarse (CO) facies. Right, image of coarse-cumulate depositional cycles. (B) CO and CU crystals, separated by a sharp transition that indicates a change in hydrological lake conditions. Peculiar cloudy Ćuid inclusion bands (FIB) in CO. (C) Common primary FIB from sample CO2-BOT. (D) CU assemblages bounded by countless single CU crystals.

Chapter 5 .

 5 Reconstructing lake bottom water temperatures and their seasonal variability in the Dead Sea Basin during MIS 5e at different temperatures to establish both an isochore (monophasic FI) and a solid-liquid-vapour equilibrium (SLVE; biphasic FI) curve(Figure 5.5). The intersection of both curves is deĄned as T x(Figure 5.6). To obtain biphasic FIs, large temperature excursions are necessary by means of an increase from room temperature to +130 °C, then back to room temperature. This causes plastic deformation of the salt host crystal with an increase in volume, forcing cavitation of a vapour bubble in the FIs. At this point, the FIs enters the stability Ąeld and represents ambient pressure conditions (Figure5.5). During the cavitation process, the chemical composition of an individual Ćuid inclusion does not change. Data are presented here as mean T x values for each sub-section (i.e., uppermost samples in coarse layers, labelled "CO-TOP" and lowermost samples in coarse layers labelled "CO-BOT").

  T x of FIs were measured in three successive CO halite layers (CO1, CO2, CO3) (Figure5.3 A) with increasing core depths (from 315.138 to 315.052 m blf, Table5.1). Additionally, the T x of FIs were measured in one Ąne CU layer (CU4) that deposited prior to the youngest CO3 halite layer (CO3-BOT) (Figure5.3 A). Here the pressure-corrected temperature (T f ), with reported uncertainties as standard error of the mean (SE), is reported for each interval. The deepest and oldest layer CO1 has been divided into three sub-sections (total number of measured FIs n = 103) with T f values of 18.4 ± 0.5 °C, 21.7 ± 0.4 °C and 21.4 ± 0.4 °C for CO1-BOT, CO1-MID and CO1-TOP, respectively. The second CO layer CO2 is divided into three sub-sections (n = 90). T f values vary from 18.3 ± 0.5 °C, 19.9 ± 0.4 °C and 22.9 ± 0.6 °C for CO2-BOT, CO2-MID and CO2-TOP, respectively (Figure 5.3 B). The uppermost and relatively youngest CO layer CO3, divided into two sub-sections, base and top (n = 48), provides T f values of 21.6 ± 0.5 °C and 23.1 ± 0.7 °C for CO3-BOT and CO3-TOP, respectively (Table 5.1, Figure 5.3 B). This dataset shows that mean T f values in each CO layer increase with increasing layer thickness from base to top. A clear difference in T f values is therefore observed by separating the dataset into groups of only coarse-bottom and coarse-top values, which reveals that lowest temperatures correspond to samples from the coarse-bottom layers, while highest temperatures are found only in coarse-top layers (Figure 5.4). Three CU sub-samples provide uncorrected T x values of 16.1 ± 1.4 °C (n = 14), 16.5± 0.7 °C (n = 26) and 15.9 ± 1.0 °C (n = 15) for CU4-A, CU4-B and CU4-C, respectively (Table5.1, Figure5.3 B). Hence, those temperatures are lower than both surrounding uncorrected T x values of CO layers with 16.8 °C and 18.1 °C for CO3-BOT and CO2-TOP, respectively(Figure 5.8). As cumulate crystals form in the water column between the surface and the bottom of the lake, the unknown temperature correction is less or equal to the temperature correction for coarse crystals. Therefore, the T f of these cumulate crystals must be less than the T f for Chapter 5. Reconstructing lake bottom water temperatures and their seasonal variability in the Dead Sea Basin during MIS 5e

Figure 5 .

 5 Figure 5.3 (A) Halite layers in core section 5017-1-A-122-2. Black lines mark facies boundaries between layers of coarse halite (CO) and cumulate halite (CU). Sampling locations of CO (orange) and CU (black) indicatedby squares. Black arrow = growth direction. (B) Corrected coarse T f data (orange squares), uncorrected cumulate T x data (black squares). Corrected cumulate T f data (triangles) show the uncertainty of BS for cumulate halite. Dashed black box marks monitored deep lake temperatures in the Dead Sea for the period 1996Ű2010 (from Gertmanand Hecht, 2002; Gertman et al., 2010). The uncertainties are reported as both, standard error of the mean (SE; error bars), and 95% conĄdence intervals (black lines). CO layers highlighted by vertical bands. (C) Monthly monitored deep lake temperatures between 2004Ű2010 (blue squares; fromGertman et al., 2010). The vertical bands illustrate the approximate period during which CO halite crystals grow on the deep lake Ćoor.

Figure 5 .

 5 Figure 5.4 (Above) Sketch illustrating the annual depositional cycles in sediment core section 5017-1-A-122-2 with sub-sample locations. (Below) Temperature offset between grouped data in terms of the interpreted growth period: deep winter cumulate halite (uncorrected T x , black squares; corrected T f , triangles), late winter/early spring coarse halite (blue squares), and late summer coarse halite (red squares). Note that the corrected cumulate data reĆect maximal potential palaeo lake temperatures, and thus should be considered with caution (see text for more details). Coloured (red and blue) envelopes are presented as the standard error of the mean. Vertical bands indicate growth season of coarse halite crystals on the deep lake Ćoor.

Figure S1 :

 S1 Figure S1: (Upper panel) Example of a typical thermal path of both monophasic (black dashed line; isochore) and biphasic (red solid line; SLV E) FIs by measuring the sound velocity at four temperature steps. The intersection of both lines (T x ) gives the entrapment temperature (T f ) if the sample has precipitated at ambient pressure. In most geological cases, a pressure-correction is needed. (Lower panel) Sketch illustrating the basic concept of Brillouin spectroscopy by using Ćuid inclusions (FIs). A monophasic FI follows a quasi-isochore (dashed black line)upon heating (red arrow) with increasing internal pressure. During cooling (blue arrow) the internal pressure becomes at some point negative, causing the FI to stretch to overcome the metastable phase (dashed black rectangles). By releasing energy during the nucleation of a vapor bubble, the biphasic FI becomes stable again. By theory, the thermal path of a biphasic FI follows the solid-liquid-vapor equilibrium (SLV E) that represents ambient pressure conditions.

Figure 5 . 5 Figure 5 . 6

 5556 Figure 5.5 S1

Figure 5 . 7

 57 Figure 5.7 Corresponding histogram to displayed single T x intersections in Supplementary Fig. 2. Bold red line marks the mean of T x values (number of measured FIs in section = 26), gray shaded area shows the standard deviation interval (SD).

Chapter 5 .

 5 Reconstructing lake bottom water temperatures and their seasonal variability in the Dead Sea Basin during MIS 5e

Figure 5 . 8 (

 58 Figure 5.8 (Left) Sampled core section (315.052 Ű 315.038 m below lake Ćoor) with coarse-grained halite layers (CO1-CO3, orange squares) and one sampled Ąne-grained cumulate layer (CU4, black squares). Black arrow indicates direction of halite growth. (Right) Mean T x values derived from 5.1 for each sub-layer. Error bars calculated with 5.3.

Chapter 6 .Figure 6 . 1

 661 Figure 6.1 Overview of global climate conditions during the past 400 ka. Left:DeĄned MIS stages plotted together with atmospheric oxygen isotope composition(Sowers et al., 1993; Jouzel et al., 1996; Malaizé et al., 1999), ice volume changes, atmospheric CO 2 variations, and Antarctic temperatures reconstructed from oxygen isotopes(Petit et al., 1999). Glacial-interglacial sub-stages of MIS 5 and 7 indicated by a to e, climatic transitions marked by Terminations I-IV. LIG: Last interglacial. Right: MIS 7 with three short-term warming sub-stages (a, c, e) and two glacial intervals (b, d), reconstructed by planktonic foraminiferal oxygen isotopes and tree pollen abundances(Tzedakis et al., 2004), and deuterium isotopes from the Vostok ice core(Petit et al., 1999). Yellow vertical bands in both plots show the halite units studied in this thesis (MIS 7c = 5017-1-A-176-2; MIS 5e = 5017-1-A-122-2).

Figure 6 . 2

 62 Figure 6.2 (A) Lithology of core 5017-1 (from Ben Dor et al., 2019; modiĄed after Neugebauer et al., 2014). (B) Age model with both core sections studied during this PhD thesis (after Goldstein et al., 2020).

Figure 6

 6 Figure 6.3 (A) Lithology of core section 176-2. Direction of deposition marked by red arrow. (B) Four segments (1-4) showing coarse-cumulate halite layering. All sub-samples are labelled next to their location (red bars). Illustrated lithology below for each segment with core section depths and numbered coarse layers. (C) Examples of CO halite growth zonations (1), three adjacent FIs within the same growth band, and a FI in monophasic (3) and biphasic (4) state.

Figure 6 . 4 Figure 6 . 5

 6465 Figure 6.4 Palaeotemperatures from core section 176-2 (MIS 7c). Blue / red points labelled with respective layer numer (see sketch above). Blue points were sampled in the lowermost parts which correspond in the upper panel to the left-handed part of each grey bar. The inverse accounts for red points. The SE of each data points is presented by the grey band.

Figure 6 . 6

 66 Figure 6.6 C. Further, the summer-winter cooling seems to be slightly pronounced (layers 12-11), indicated by ∆T cooling values ≤ 0 °C. As Ągured out, the Ąrst post-meromictic year 0 (= 1996 CE) is associated with a warming by +1 °C (∆T warming: difference of maximal and minimal monitored deep lake water temperatures). A similar but less strong increase can be seen by ∆T cooling (<1 °C), thereby both parameters practically describe the annual seasonality. ∆T cooling never exceeds ∆T warming during the post-meromictic years in the modern DS until year 4 (= 2000 CE) (Figure 6.6 A). Interestingly, both curves of ∆T show parallel decreasing and increasing trends from years 2 to 4 (=[START_REF] Yechieli | Will the Dead Sea die?[END_REF] -2000 CE) CE), respectively, and reach their highest rates when the late-summer lake temperature rises more pronounced by exceeding 24 °C. The derived estimates of T f are in line with modern deep lake temperatures in the Dead Sea which ranged between 22-25 °C between 1996-2010 after the onset of holomictic conditions in 1996, coeval with the nucleation of halite(Gertman et al., 2010). The averaged increase in lake-bottom temperatures can be estimated to +0.38 °C per year which is more than the actual DS surface warming (+0.8 °C per decade)(Kishcha et al., 2021). Averaged temperature increases for each season speciĄcally yield values of both +0.3 °C/year during late winter, and +0.18 °C/year for late-summer episodes, respectively.

Figure 6 . 7

 67 Figure 6.7 Characteristics of core section 176-2. A) Thickness of warming/cooling varves. B) Warming varve thickness vs. cumulate halite layer thickness. C) Cooling/warming varve thickness of each layer vs. cumulate halite layer thickness of each varve. D) Seasonality of summer-winter (blue) and winter-summer (red), respectively.

Figure 6 . 8

 68 Figure 6.8 Corrected Ćuid inclusion temperatures in core 5017-1-A-176-2. The yellow band shows the present-day mean of DS deep waters (including an error margin of ± 0.6 °C).The blue band displays the mean of corrected Ćuid inclusion data from core 5017-1-A-122-2 (MIS 5e) (error margin of ± 0.5 °C), the purple band marks the mean of core section 176-2. Note that the error overlaps with the gray band due to almost identical values.

Chapter 6 .Figure 6 .

 66 Figure 6.10 (A) Sketch of tested varve deĄnitions used in this study. Note that both warming and cooling varve distances enclose a complete layer. (B) Terminology for both halite types in core 176-2 along one growth cycle (n -> n+1). (C) Calculation of relative temperature differences of a data set for both layer n and n+1, respectively. Black bars symbolise the location of sub-samples.

  intervals of I low (CO13 to CO8) yield constant values of ∆V top = 4.5 cm (Figure 6.11 C). The transition towards I cen shows remarkably increases of relative thicknesses to 6.2 cm and 7.3 cm (CO9-TOP to CO8-TOP, and CO8-TOP to CO7-TOP, respectively). Subsequently, a declining trend follows with values of 5.0 cm, 2.5 cm, 3.5 cm, and 2.8 cm , with the two latter values being part of I upp . From CO3-TOP to CO2-TOP, ∆V top increases by 1.0 cm, while the last sub-interval (CO2-TOP to CO1-TOP) is marked by a sharp increase to ∆V top = 8.4 cm. As a main result, ∆V top is in phase with T x values of CO-TOP sub-samples along the whole core section (Figure 6.11 D).

  thickness late winter cycle (a) Relative varve thicknesses measured from CO-BOT(n) to CO-BOT(n+1).

  T x bottom data plotted with curve from (a).

  thickness [cm] Relat. varve thickness late summer cycle (c) Relative varve thicknesses measured from CO-TOP(n) to CO-TOP(n+1).

  thickness [cm] Relat. varve thickness late summer cycle (d) T x from coarse top data plotted with curve from (c).

Figure 6 .

 6 Figure 6.11 Combined results of core section 176-2. Standard error indicated by blue (b) and red (d) bands surrounding the mean T x values.

Figure 7 . 1

 71 Figure 7.1 Gypsum-transformation to the hemi-hydrate phase after heating to 130 °C for ca 60 seconds.

Figure 7 . 3

 73 Figure 7.3 Lab-grown gypsum with trapped FIs. Black bubble mark the biphasic phase of FIs. A, B, D-G) are examples to show the various shapes observed during the experiments. C) White areas are indicative for gaseous inclusions without any liquid. Scale bars are equally to 2 µm in D), 13 µm in F, and 6 µm in G), respectively.

Figure 7 . 4

 74 Figure 7.4 Lab-grown gypsum data from Ćuid inclusions from three evaporation experiments. A-C: Dissolved CaSO 4 in pure H 2 O and evaporated at 21 °C. D: Dissolved NaCl and CaSO 4 in pure H 2 O and evaporated at 25 °C in a water bath. E-H: Same parent solution as in D, but evaporated at 21 °C. Fitted lines represent monophasic isochores (black solid) and biphasic LVE (red dashed).

Figure 7 . 5

 75 Figure 7.5 Exemplary Ąts of Ćuid inclusion data of natural gypsum samples from Table C.1. Fitted lines of monophasic FIs (isochores, black solid) and biphasic FIs (LVE, red dashed) deĄne the T x intersection. See Appendix A for additional Ąts of other samples listed in Table C.1.

  Second, lab-grown gypsum data showed often abnormal Ątted curves of isochores and/or SLVE (Figure7.4 B, C, F), or even elevated SLVE data points, leading to strongly deviated entrapment temperatures, e.g., 16.3 °C instead of 25 °C, or 46 °C instead of 21 °C (Figure7.4 D, H). Nevertheless, natural gypsum samples that are much larger with higher abundances of trapped FIs seem to be suitable to determine T x values (Figure7.5). The main contrast to lab-grown gypsum data is the linear Ątting of isochores with relatively

Figure 8 . 1

 81 Figure 8.1 Calculated and measured sound velocity data for pure water.

Figure 8 . 3

 83 Figure 8.3 Temperature distribution of w values decreases with increasing NaCl concentration.

Chapter 8 .Figure 8 . 5 Figure 8 . 6

 88586 Figure 8.5 Comparison of sound velocity data of water samples from the Aral Sea. Groundwater pounds (Aral11, Aral13) vs. bottom waters and surface waters.

Chapter 8 .Figure 8 .

 88 Figure 8.10 Observed Ąnal ice melting points as a function of measured sound velocity data of biphasic FIs in sample MG1. A: Crystal 05, B: crystal 09, C: crystal 07, D: all data grouped by their respective colors from A-C. Pure water indicated at y = 0 with w = 1496 m/s by squared box.

8. 5 Figure 8 .

 58 Figure 8.11 Calculated salinity from ice melting points in biphasic gypsum FIs vs. measured sound velocity. Plots presented for sub-samples 05 (A), 09 (B), 07 (C), and a complete plot with data from A-C.

Chapter 8 .Figure 8 .

 88 Figure 8.13 Final plot of salinity vs. sound velocity data from biphasic gypsum samples

2. 3

 3 °C in summers. The results are in general agreement with a short-term global warm phase, which has already been proven by other proxy methods. In relation to the current warm phase and rising global temperatures, especially in the Middle East, the results show that the region and the Dead Sea are very sensitive to external regional or global warm phases. In the future, further studies should be carried out on interglacial salt layers in drill core 5017-A to create a uniform data set for comparing deep water temperatures.The application of Brillouin spectroscopy was further tested on laboratory-grown gypsum samples by comparing measured entrapment temperatures with known growth temperatures. The work proved challenging due to an insufficient amount of trapped FIs in gypsum crystals. However, a few temperatures could be obtained and showed the potential for future research on gypsum samples containing a larger amount of FIs. Nevertheless, it was shown that natural gypsum samples could be suitable for palaeoclimatic purposes, although the data obtained might correspond to secondary FIs.However, an encouraging point that inspires us to continue the methodological development on gypsum is that, data from biphasic FIs in natural gypsum indicate a positive correlation between sound velocity and increasing salinity. Thus, sound velocity can be used as a salinity proxy in evaporites what is also supported by data from capillary solution experiments with increasing NaCl concentrations (0 -26.4 wt. % NaCl). Thus, sound velocity measurements of biphasic FIs in natural gypsum samples might be a promising tool for future palaeosalinity reconstructions.

(

  reduce the voltage manually to 0 V, then click "reĆected light off"). Wear your GOGGLES before you turn the shutter. In reĆected light mode, you need to turn a silver wheel right next to the microscope light exit to unblock the light.Otherwise you will have no laser signal. In order to get the highest signal, check and compare the spectrum at several depths while turning the big knob at the microscope control. If you use a capillary, it might be necessary to go deeper manually. Stop observing and click "start acquisition" to detect the spectrum. 13. Filename structure for Mathematica analysis package: Sample name_number of crystal_FI number_number of FI measurement_L_ (liquid) or _LV (liquid-vapor) _FI size_FI depth_objective_pinhole_number of scans_time_temperature.DAT Chapter B. Ruling out thermal effects due to the laser beam measured intensity by using IF old (all in mW). For instance, if approximately doubling the intensity of a current setting with 54 mW on a sample and measured 144 mW, B.3 yields: a value of around 269 mW must be reached on a powermeter to increase the intensity on a sample to 101 mW. A proportional ratio of 3higher (new) intensity after the Ąlter and b the lower (older) intensity after the Ąlter in mW. To keep a standard baseline of ca. 300 counts, we used B.6: b a * b i = baseline (B.6) with b i the intensity on the sample in mW. For an increase in laser intensity, Following equation B.6, the new (reduced) intensities on a sample are: , no signiĄcant increase in speed of sound values was observed. T x data showed similar values within the instrumental error for one FI, by using both 54 mW and 101 mW for its isochores and LVE measurements. Based on these Ąndings, an induced heating of the sample by increased laser intensity can be ruled out. B.2 Results experiment A Initially, a cubic-shaped halite Ćuid inclusion (FI01Hal24; length = 18 µm, width = 15 µm, depth = 36 µm) was selected from a controlled growth experiment (Hal24 with T f = 24.5 ± 1 °C) for Brillouin analysis. Isochores of FI01Hal24 were measured at 4 descending temperature steps (29, 21, 13, 5 °C) at 101 mW.Then, the setting was changed to 28 mW and isochores were measured at same of scans (and thus the acquisition time) to reach 300 counts are lowest at 101 mW and increase with lowering the intensity to 54 and 28 mW, respectively. The Ątted curve for 3 monophasic FIs at 4 temperatures(29, 21, 13, 5 °C) shows no difference in sound velocity values at varying intensity, with a typical range of < 15 m/s (Figure B.4). In contrast, Ątted SLVE curves vary stronger between 54 and 101 mW and show larger differences between both SLVE curves with increasing temperature (Figure B.5). Although the number of FIs (n < 4) is too low to determine the most accurate entrapment temperature, one can still compare the mean T x values between varying laser intensity. The lowest intensity at 28 mW yields a mean T x of 21.8 ± 1.1 °C, the intermediate intensity at 54 mW gives 24.3 ± 0.4 °C and the highest intensity has a mean T x value of 24.4 ± 1.8 °C (Table B.1).

  T x measurements. The T x data for sample Hal24 reveal a mean value of 24 ± 1.6 °C (n = 24), which is in good agreement with the known formation temperature of the host NaCl crystal of 24.5 ± 1 °C. Further, 13 T x values are within the error range of 1 °C during NaCl crystallization (23.5 < T x < 25.5 °C).

  Figure B.3

  2 + or Sr 2 +, forming molecules by mainly covalent

	or ionic bounds with anion-complexes, e.g., CO 3	2-(carbonate group), SO 4	2-

(sulfate group), or Cl -(chloride group)

  On the contrary, the sulphate mineral gypsum (CaSO 4 • 2 H 2 O) has an extremely low solubility of 2.4 g/L at 25 °C (International Labour Organization, n.d.), deĄning a narrow range of possible amounts of dissolved/precipitated gypsum of a brine.

  1.3 Fluid inclusionsWhile a crystal nucleates from a supersaturated brine, solution, or Ćuid, crystal defects in the lattice are Ąlled by the surrounding supersaturated liquid or gas, leading to local cavities. The occurrence of local cavities in solid crystals is known for approximately 230 years since the Ąrst observation of petroleum-Ąlled cavities in quartz

  ).

	Chapter 1. Introduction
	For example, if NaCl with a solubility of maximal 360 g/L is dissolved in 1 L of
	H 2 O, it has a volume percentage of 36%, but a mass percentage of 360 g/(1000
	g H 2 O + 360 g) x 100 = 26.47 weight % NaCl. Therefore, in Ćuid inclusions studies
	of geological interest, the traditional expression of salinity is reported in weight %

  Menzel, ø 16 mm) immediately afterwards, both openings were sealed with a slightly heated blue wax. The same procedure was applied on other solutions with different ionic compositions (pure H 2 O, CaSO 4 + pure H 2 O, CaSO 4 + NaCl + pure H 2 O).

). For each sample, a rectangular borosilicate capillary (Vitrotubes T M ; 0.05 x 0.5 mm) was dipped into the solution within seconds until it was about three quarters Ąlled. Placed on a borosilicate coverslip (

Table 4 . 1

 41 Published sound velocity values for increasing NaCl concentrations at 25 °C (fromChen and Millero, 1986).

	Molality b Weight % NaCl w 0 N aCl,relative [mol/kg] [m/s] [m/s]	w 0 N aCl,absolute
	0.2519	1.5	15.88
	0.4906	2.8	30.14
	0.7541	4.2	46
	1.0399	5.7	62.14
	1.5935	8.5	93.46
	2.3418	12.0	132.14
	3.3723	16.5	183.24
	3.8718	18.5	206.17
	4.5985	21.2	237.96
	4.8797	22.2	249.32
	5.0751	22.9	256.98
	5.4347	24.1	271.04
	6.1469	26.4	298.17

Table 4 .2 Properties

 4

	Mass volume H 2 O	S	m NaCl Moles of solute Molality b
	[g]	[wt.% NaCl]	[g]	[mol]	[mol/kg]
	50	0.5	0.2492	0.0043	0.086
	50	2.5	1.2793	0.0214	0.439
	50	5	2.6299	0.0428	0.901
	50	10	5.5604	0.0856	1.901
	50	15	8.8198	0.1283	3.020
	50	20	12.5004	0.1711	4.278
	50	25	16.6710	0.2139	5.704
	50	26.4	17.9291	0.2267	6.167

of prepared NaCl capillary samples. S = Salinity, m = Mass.

Reconstructing lake bottom water temperatures and their seasonal variability in the Dead Sea Basin during MIS 5e

  Chapter 5. Reconstructing lake bottom water temperatures and their seasonal variability in the Dead Sea Basin during MIS 5eAbstractInterglacial periods are characterised by thick accumulations of halite units in the Dead Sea Basin. During these intervals, small water droplets (Ćuid inclusions)

This chapter was published in the journal ŠThe Depositional RecordŠ in February 2022. https://doi.org/10.1002/dep2.185 Authors: Brall, N., Gardien, V.,

Ariztegui, D., Sorrel, P., Guillerm, E. and Caupin, F.

Table 5 . 1

 51 Measured entrapment temperatures in halite Ćuid inclusions. Columns provide the sediment core depth in metres below lake Ćoor (m blf), the number of Ćuid inclusions per sample (N), the averaged entrapment temperature per sample (Mean T x ), the standard deviation of a sample population (SD), and the standard error of the mean (SE). Both lower and upper boundaries of 95% conĄdence intervals of each sample mean calculated with studentŠs t-test. Coarse sampled were corrected for hydrostatic pressure (Mean T f ) (see Methods). Mean T f values were not determined for cumulate halite samples (N.A.).

	Sample	Halite Depth Section N Mean SD SE Lower Upper Mean Lower Upper
	name	facies in core depth	T x	95% 95% T f 95% 95%
		5017-1			CI	CI	CI	CI
		(m blf) (cm)	(°C) (°C) (°C) (°C) (°C) (°C) (°C) (°C)
	CO3-TOP coarse 315.052 2.2 31 18.3 3.9 0.7 16.9 19.7 23.1 21.7 24.5
	CO3-BOT coarse 315.055 2.5 17 16.8 2.1 0.5 15.7 17.9 21.6 20.5 22.7
	CU4-C cumulate 315.074 4.4 15 15.9 3.9 1 13.7 18.1 N.A. N.A. N.A.
	CU4-B cumulate 315.09	6	26 16.5 3.4 0.7 15.1 17.9 N.A. N.A. N.A.
	CU4-A cumulate 315.095 6.5 14 16.1 5.2 1.4 13	19.1 N.A. N.A. N.A.
	CO2-TOP coarse 315.103 7.3 21 18.1 2.6 0.6 16.9 19.3 22.9 21.7 24.1
	CO2-MID coarse 315.105 7.5 35 15.1 2.4 0.4 14.3 15.9 19.9 19.1 20.7
	CO2-BOT coarse 315.109 7.9 34 13.5 3 0.5 12.4 14.5 18.3 17.2 19.3
	CO1-TOP coarse 315.12	9	37 16.6 2.1 0.4 16	17.4 21.4 20.8 22.2
	CO1-MID coarse 315.124 9.4 40 16.9 2.6 0.4 16	17.7 21.7 20.8 22.5
	CO1-BOT coarse 315.138 10.8 26 13.6 2.5 0.5 12.6 14.6 18.4 17.4 19.4

Table 6 . 1

 61 Measured uppermost (TOP) and lowermost (BOT) sub-intervals in coarse halite layers in core 5017-1-A-176-2.The central interval I cen describes the layers from CO7-BOT to CO4-BOT (445.37 to 445.255 m blf; 11.5 cm length). The maximum of ∆ T warming along the whole core (+3.2 °C) is observed in layer CO7 by a sharp warming along the 2-cm thick coarse layer. This peak is followed by a stepwise two-year decline in both CO-BOT

	Sample	Section 5017-1	T x	N 1 SD 2 SE 3	CI 4	-z * 5	+z * 6
	name	depth	depth mean				95%
		[cm]	[m blf]	[°C]		[°C] [°C] [°C]	[°C]	[°C]
	CO1-TOP	6	445.09	19.8	31	2.9	0.5	1.02	18.8	20.8
	CO1-BOT	11.8	445.148 18.7	25	2.9	0.6	1.14	17.6	19.8
	CO2-TOP	14.4	445.174 20.2	28	2.9	0.5	1.07	19.1	21.3
	CO2-BOT	15.7	445.187	20	14	2.5	0.7	1.31	18.7	21.3
	CO3-TOP	18.2	445.212 18.7	22	2.6	0.6	1.09	17.6	19.8
	CO3-BOT	20	445.23	17.5	26	3	0.6	1.15	16.3	18.7
	CO4-TOP	21	445.24	18.5	29	3	0.6	1.09	17.4	19.6
	CO4-BOT	22.5	445.255 16.2	16	2.3	0.6	1.13	15.1	17.3
	CO5-TOP	24.5	445.275 17.8	22	2.8	0.6	1.17	16.6	19.0
	CO5-BOT	25.5	445.285 16.5	21	2.4	0.5	1.03	15.5	17.5
	CO6-TOP	27	445.3	18.8	22	2.4	0.5	1.00	17.8	19.8
	CO6-BOT	28.5	445.315 17.6	16	3	0.7	1.47	16.1	19.1
	CO7-TOP	32	445.35	19.4	20	1.6	0.4	0.70	18.7	20.1
	CO7-BOT	34	445.37	16.2	21	2.5	0.5	1.07	15.1	17.3
	CO8-TOP	39.3	445.423 16.8	15	2.7	0.7	1.37	15.4	18.2
	CO8-BOT	42	445.45	18.2	19	2.5	0.6	1.12	17.1	19.3
	CO9-TOP	45.5	445.485 16.9	24	2.7	0.6	1.08	15.8	18.0
	CO9-BOT	47	445.5	16	17	2.4	0.6	1.14	14.9	17.1
	CO10-TOP	50	445.53	17.5	17	1.8	0.4	0.86	16.6	18.4
	CO10-BOT	51.5	445.545 15.5	14	2	0.6	1.05	14.5	16.5
	CO11-TOP	54.5	445.575	17	25	2.5	0.5	0.98	16.0	18.0
	CO11-BOT	57	445.6	17.1	13	1.4	0.4	0.76	16.3	17.9
	CO12-TOP	59	445.62	16.8	19	2.7	0.6	1.21	15.6	18.0
	CO12-BOT	60	445.63	17.4	16	2.2	0.5	1.08	16.3	18.5
	CO13-TOP	63.5	445.665 17.5	18	2.6	0.6	1.20	16.3	18.7
	CO13-BOT	66.5	445.695 14.8	15	2.8	0.7	1.42	13.4	16.2
	Total	60.5 7	0.605 7 17.6 8 525 2.5 8 0.6 8 1.10 8 16.5 8 18.7 8
	1 number of measured Ćuid inclusions per sample and in total
	2 standard deviation of a sample population
	3 standard error of the mean
	4 conĄdence interval
	5 T x mean -CI 95%

Table 6 . 2

 62 Pressure-corrected T f data of core section core 5017-1-A-176-2.

	Sample	Section 5017-1	T x	N
	name	depth	depth mean
		[cm]	[m blf]	[°C]
	CO1-TOP	6	445.09	22.7 31
	CO1-BOT	11.8	445.148 21.6 25
	CO2-TOP	14.4	445.174 23.1 28
	CO2-BOT	15.7	445.187 22.9 14
	CO3-TOP	18.2	445.212 21.6 22
	CO3-BOT	20	445.23	20.4 26
	CO4-TOP	21	445.24	21.4 29
	CO4-BOT	22.5	445.255 19.1 16
	CO5-TOP	24.5	445.275 20.7 22
	CO5-BOT	25.5	445.285 19.4 21
	CO6-TOP	27	445.3	21.7 22
	CO6-BOT	28.5	445.315 20.5
	CO7-TOP	32	445.35	22.3 20
	CO7-BOT	34	445.37	19.1 21
	CO8-TOP	39.3	445.423 19.7 15
	CO8-BOT	42	445.45	21.1 19
	CO9-TOP	45.5	445.485 19.8 24
	CO9-BOT	47	445.5	18.9 17
	CO10-TOP	50	445.53	20.4 17
	CO10-BOT	51.5	445.545 18.4 14
	CO11-TOP	54.5	445.575 19.9 25
	CO11-BOT	57	445.6	20.0 13
	CO12-TOP	59	445.62	19.7 19
	CO12-BOT	60	445.63	20.3 16
	CO13-TOP	63.5	445.665 20.4 18
	CO13-BOT	66.5	445.695 17.7 15

Table 6 .4

 6 Measured lowermost sub-intervals (BOT) in coarse halite layers in core 5017-1-A-176-2.

	Sample	Section 5017-1	T x	N SD SE	CI	-z *	+z *
	name	depth	depth mean				95%
		[cm]	[m blf]	[°C]		[°C] [°C] [°C] [°C] [°C]
	CO1-BOT	11.8	445.148 18.7	25	2.9 0.6 1.14 17.6 19.8
	CO2-BOT	15.7	445.187	20	14	2.5 0.7 1.31 18.7 21.3
	CO3-BOT	20	445.23	17.5	26	3	0.6 1.15 16.3 18.7
	CO4-BOT	22.5	445.255 16.2	16	2.3 0.6 1.13 15.1 17.3
	CO5-BOT	25.5	445.285 16.5	21	2.4 0.5 1.03 15.5 17.5
	CO6-BOT	28.5	445.315 17.6	16	3	0.7 1.47 16.1 19.1
	CO7-BOT	34	445.37	16.2	21	2.5 0.5 1.07 15.1 17.3
	CO8-BOT	42	445.45	18.2	19	2.5 0.6 1.12 17.1 19.3
	CO9-BOT	47	445.5	16	17	2.4 0.6 1.14 14.9 17.1
	CO10-BOT	51.5	445.545 15.5	14	2	0.6 1.05 14.5 16.5
	CO11-BOT	57	445.6	17.1	13	1.4 0.4 0.76 16.3 17.9
	CO12-BOT	60	445.63	17.4	16	2.2 0.5 1.08 16.3 18.5
	CO13-BOT	66.5	445.695 14.8	15	2.8 0.7 1.42 13.4 16.2
	Total	54.7	0.547 17.1 233 2.5 0.6 1.14 15.9 18.2

Table 6 . 5

 65 Splitted part of the Ąrst 26 cm in core section 176-2 with uppermost sub-intervals (TOP) in coarse halite layers.

	Sample	Section 5017-1	T x	N	SD SE	CI	-z *	+z *
	name	depth	depth mean				95%
		[cm]	[m blf]	[°C]		[°C] [°C] [°C] [°C] [°C]
	CO1-TOP	6	445.09	19.8	31	2.9 0.5 1.02 18.8 20.8
	CO2-TOP	14.4	445.174 20.2	28	2.9 0.5 1.07 19.1 21.3
	CO3-TOP	18.2	445.212 18.7	22	2.6 0.6 1.09 17.6 19.8
	CO4-TOP	21	445.24	18.5	29	3	0.6 1.09 17.4 19.6
	CO5-TOP	24.5	445.275 17.8	22	2.8 0.6 1.17 16.6 19.0
	CO6-TOP	27	445.3	18.8	22	2.4 0.5 1.00 17.8 19.8
	CO7-TOP	32	445.35	19.4	20	1.6 0.4 0.70 18.7 20.1
	Total	26	0.26	19.0 174 2.6 0.5 1.02 18.0 20.0

Table 6 . 6

 66 Splitted part of the lowest 24 cm in core section 176-2 with uppermost sub-intervals (TOP) in coarse halite layers.

	Sample	Section 5017-1	T x	N	SD SE	CI	-z *	+z *
	name	depth	depth mean			95%
		[cm]	[m blf]	[°C]		[°C] [°C] [°C] [°C] [°C]
	CO8-TOP	39.3	445.423 16.8	15	2.7 0.7 1.37 15.4 18.2
	CO9-TOP	45.5	445.485 16.9	24	2.7 0.6 1.08 15.8 18.0
	CO10-TOP	50	445.53	17.5	17	1.8 0.4 0.86 16.6 18.4
	CO11-TOP	54.5	445.575	17	25	2.5 0.5 0.98 16.0 18.0
	CO12-TOP	59	445.62	16.8	19	2.7 0.6 1.21 15.6 18.0
	CO13-TOP	63.5	445.665 17.5	18	2.6 0.6 1.20 16.3 18.7
	Total	24.2	0.242 17.1 118 2.5 0.6 1.12 16.0 18.2

Table 6 .7

 6 Compilation of measured varve thicknesses and calculated seasonal temperatures as deĄned in Figure6.10 for core section 176-2.

	∆ T ∆ T	cooling warming
	Layer Relative Relative Coarse Cumulate Cooling Warming	layer varve varve number CO-TOP CO-BOT layer	[cm]

Table 7 . 1

 71 Overview on measured temperatures of lab-grown gypsum samples compared with the known growth temperatures.

	Type	Sample name	FI name T x	Growth	Offset
				[°C] temperature [°C]	[°C]
	Lab-grown SG032020-210602	1	22.0	21	-1.02
	Lab-grown SG032020-210602	9	22.4	21	-1.35
	Lab-grown SG032020-210602	11	18.6	21	2.36
	Lab-grown SGH21_22-191217	1	21.4	21	-0.43
	Lab-grown SGH21_22-191217	2	46.0	21	-25.01
	Lab-grown SGH21-400-191119	1	22.0	21	-0.99
	Lab-grown SGH21-400-191119	2	20.0	21	1.04
	Lab-grown SGH25-2_01-FI02	2	16.3	25	8.69

  4 • 2 H 2 O has affected the crystal sizes of lab-grown samples, which caused low entrapment abundances and/or only tiny FIs that were trapped. Another point that was observed is the formation of larger water assemblages within the Ćat gypsum crystals which did not show the classical behaviour of FIs. Upon cooling-heating, these water assemblages changed their location, shape, and volume by incorporating adjacent FIs. Therefore, the number of T x values (= 8) obtained from lab-grown gypsum samples was quite small and needed two years of lab work. For comparison, within one day of lab work it is normally possible to obtain 20 T x values from a single lab-grown NaCl crystal.

Table 8 .1 Measured

 8 H 2 O data are generally speaking ≥ 2.3% lower than literature values on the 5-65 °C temperature range (Table 8.1; Figure 8.1). This error, expressed in % from the Ąnal calculated value, Chapter 8. Towards a new proxy for palaeosalinity: Sound velocity of saline aqueous solutions and Ćuid inclusions in gypsum

	rises with increasing temperatures to 2.76%. For ambient moderate temperatures
	(15-35 °C), the error does not increase signiĄcantly during a 20 °C shift (0.1%),
	thus it is lower than the instrumental error of 0.15% (3 m/s). Such a shift has been
	expected due to the refraction of light at interfaces as mentioned in 4.1. Hence, a
	capillary correction factor (c) needs to be adapted for all capillary samples which
	is expressed by the difference to literature data and further related to the fraction
	of measured:calculated values.

8.2 Saline aqueous solutions 8.2.1 Pure water On a total temperature range of 60 °C (5-65 °C), the measured w H 2 O data vary between 1393 to 1510 m/s (∆117 m/s; Table 8.1). At 25 °C, w H 2 O = 1457 m/s, thus, 40 m/s less than the literature value of 1497 m/s. Pure w The obtained w data for H 2 O are close to those from the literature with a maximum error of 1.2 % at 5 °C (Table 8.1). The Ątted data are best described by a third

  O calc. ) = 1399 + 5.248x -5.798 * 10 -2 x 2 + 2.128 * 10 -4 + x 3

		1600
	Sound velocity [m/s]	1400 1500	Offset Correction Error [°C] [m/s] [m/s] [m/s] factor % 5 1444 1426 18 1.012 1.2 15 1480 1466 14 1.010 1.0 25 1509 1497 12 1.008 0.8 35 1532 1520 12 1.008 0.8
			45 1548 1536	12	1.008	0.8
	55 1558 1547 65 1564 1553 order polynomial Ąt (Figure 8.1): 15 25 Temperature [°C] 11 1.007 11 1.007 35 45 w(H 2 5 Measured Calculated	55	0.7 0.7	65

(R 2 = 1) w(H 2 O meas. ) = 1369 + 4.679x -4.623 * 10 -2 x 2 + 1.192 * 10 -4 x 3 (R 2 = 0.9999) (8.1)

The observed difference between measured and calculated w values further helps to treat data of saline samples with caution during the interpretation. Therefore, data obtained from salty aqueous solutions have been corrected by 1/1.008, while natural water samples are reported as raw output values of the analysis program because the refractive index of individual samples could not be determined. aqueous solutions and Ćuid inclusions in gypsum

Table 8 . 2

 82 Measured sound velocity values using BS at increasing NaCl concentrations at 25 °C, compared with literature w data

  Chapter 8. Towards a new proxy for palaeosalinity: Sound velocity of saline aqueous solutions and Ćuid inclusions in gypsum Moreover, the results of NaCl solutions reveal that one "working temperature" (e.g., 25 °C) is sufficient to characterise the approximate salinity regime of liquid samples as illustrated by Figures 8.2 and 8.3. Data of NaCl solutions needed to be adjusted for the refractive index n of pure water, that of NaCl solutions, and Ąnally to be corrected for the capillary effect by

	multiplying the adjusted values with 1 over the average capillary factor obtained
	from pure water:	w adj. =	w raw * n N aCl(T,msat(T )) n N aCl(T,m)	(8.2)
	with w raw the measured values by BS, n N aCl(T,msat(T )) the refractive index of
	saturated NaCl at any temperature, and n N aCl(T,m) the refractive index of NaCl at
	both any temperature and molality, respectively. It follows:
				w corrected = w adj. *	1 1.008	(8.3)
		1800
		1700
	w (m/s)	1600						5.5 15.2 °C °C 24.9 °C 34.6 °C
								44.3	°C
		1500						54.0 63.7	°C °C
		0 1400	1	2	3		4	5	.2% 6
		0.5 2.5	0.086 0.4388	1502 1524 Molality (mol/kg) 1502 1525	0 -1	1.000 1.001	0 -0.1
		5	0.901	1551	1549	2	0.999	0.1
		10	1.901	1606	1609	-3	1.002	-0.2
		20	4.278	1721	1708	13	0.992	0.8
		26.4	6.167	1795	1774	21	0.988	1.2

m/s. The high-saline capillary solution (20 wt.% NaCl) falsl between 1687<w<1710 m/s, and the saturated NaCl sample Ąnally yields a nearly constant interval of 1783<w<1791 m/s (Figure

8

.3).

Following this, a decreasing trend of w ranges coincides with increasing NaCl

  who used an extended approach of the standard UNESCO equation(IOC, 2010). Corrected NaCl capillary data with Ąts from the literature. Molality in mol/kg shown by added numbers right of the plot. Chapter 8. Towards a new proxy for palaeosalinity: Sound velocity of saline aqueous solutions and Ćuid inclusions in gypsum 8.2.3 Water samples from saline environments The results reported in this section are reported as raw w values only. Distributions of capillary w data of waters, brines, and other solutions measured by BS are presented in Figure 8.6. Three major w intervals as a function of dissolved ions and/or the sampling locations are derived from the data set: Solutions of pure H 2 O, The aforementioned solutions are followed by elevated w values of two water samples collected from the Aral Sea surface ("AS surface") and from the bottom at ca 30 m depth ("AS bottom"). The border to more than 1800 m/s is marked by a saturated NaCl sample ("NaCl sat."). The third interval deĄnes a range between 1800-2000 m/s. It contains data sets of four water samples that were collected from high-saline environments, in ascending order of w values:

	Chapter 8. Towards a new proxy for palaeosalinity: Sound velocity of saline
		aqueous solutions and Ćuid inclusions in gypsum
	1800	6.138
		4.278
	1700
	1.901 0.901 0.439 0.086 m"), and of Mediterranean seawater from a shoreline NW of Sicily ("MS Sicily") 0 10 20 30 40 50 60 1400 1500 1600 Temperature (°C) w (m/s) yield w values >1400 and <1600 m/s. Secondly, an interval between 1600-1800 m/s deĄnes intermediary solutions with, e.g., mixed NaCl, CaSO 4 compounds ("SGH25 Figure 8.4 pure CaSO 4 ("CaSO4 sat", "SG032020"), 0.01 m sodic-sulphatic ("NaCl-CaSO4 0.01 BULK", "SGH25 REP").

Table 8 .4

 8 Statistics with sound velocity and Ąnal ice melting points of gypsum MG1 sub-samples (Tunisia) Chapter 8. Towards a new proxy for palaeosalinity: Sound velocity of saline aqueous solutions and Ćuid inclusions in gypsum

		Sub-sample
		05	09	07
	mean w LVE [m/s] 1722 1707 1848
	n	23	20	17
	SD w LVE [m/s]	77	41.4 38.0
	range w LVE [m/s] 304	143	163
	min w LVE [m/s] 1550 1630 1754
	max w LVE [m/s] 1854 1773 1917
	T ice mean [°C]	-19.6 -23.9 -30.2
	n	23	21	18
	T ice SD [°C]	6.6	6.0	2.7
	T ice range [°C]	25.3 22.7 11.7
	min T ice [°C]	-32.4 -35.3 -35.7
	max T ice [°C]	-7.1 -12.6 -24.0

Table B . 1

 B1 Experiment B: Comparison T x data of halite FIs at different laser intensities.

	Laser intensity [mW] FI01 FI02 FI03 FI04 Mean T x [°C] SD [°C]
	28	22.5 20.5 22.4		21.8	1.1
	54	24.8	24	24	24.5	24.3	0.4
	101	23.1	27	24.3 23.2	24.4	1.8
	Results for NaCl

Table C . 1

 C1 Overview on T x data acquired in natural gypsum samples

	Chapter D. MG1 Ćuid inclusion data tables
	Age (Location)	Sample name	FI name	Tx [°C]
	Messinian (Eraclea Minoa, Sicily)	EM-TOP		12.3 / 37.4
	"	EM-MID		30.5
	"	EM-BOT		21.7
	"	EM-BOT		34.8
	"	EM-BOT		27.8
	Messinian (Santa Elisabetta, Sicily)	SE_cyano		20.9
	"	SE_cyano		21.5
	"	SE_cyano		17.6
	"	SE_cyano		23.7
	Holocene (Aral Sea shore)	AS14-210719		43.4
	"	AS14-210719	17	50.7
	"	AS14-210719	21	21.1
	"	AS14-210719		30.7
	"	AS14-210719		28.4
	"	AS14-210719	20	21.1
	Messinian (Piedmont, Banengo, Italy)	BAN_200716		51.9
	"	BAN_200716		29.2
	"	BAN_200716	13	36.6
	"	BAN_200716	14	35.5
	"	BAN_200716	20	36.5
	Holocene (Sabkha El Melah, Tunisia) MG1_01-201215		21.6
	"	MG1_05-210720		54.7
	"	MG1_05-210720		44.8
	"	MG1_05-210720		48.1
	"	MG1_05-210720		41.3
	"	MG1_05-210720		75.2
	"	MG1_07-210428		30.0
	"	MG1_07-210428	44	59.7
	"	MG1_07-210428	47	48.6
	"	MG1_07-210428	51	47.4
	"	MG1_07-210428	56	17.7
	"	MG1_07-210428	66	46.4
	"	MG1_07-210428	71	13.7
	"	MG1_07-210428	73	47.2
	"	MG1_08-210429		33.9
	"	MG1_10-210721	15	37.8

Table D . 1

 D1 Sub-sample 05: Fluid inclusion data from microthermometry and Brillouin Spectroscopy with calculated salinity values afterBodnar (1993).

		FI	T eu T ice w biphasic	Salinity
			[°C]	[°C]	[m/s]	[wt. % NaCl eq.]
		2	-58 -19.9	1780	22.3
		7	-56 -13.7	1649	17.5
		9	-49 -16.3	1678	19.7
		15	-49.8 -15.8	1681	19.3
		18	-52	-17	1686	20.2
		20		-21	1702	23.0
		21	-54.8 -16.5	1676	19.8
		22		-29	1762	28.0
		28	-58 -18.7	1700	21.5
		31		-23.8	1726	24.8
		34		-27	1848	26.8
		50		-32.4	1854	30.2
		52	-55.4 -17.7	1746	20.7
		54		-25.2	1745	25.7
		55		-26	1814	26.2
		104		-20	1689	22.4
		150 -54.8 -26.6	1752	26.6
		151		-24.4	1833	25.2
		152		-10.3	1623	14.3
		153		-15.3	1689	18.9
		154		-7.4	1630	11.0
		155		-7.1	1550	10.6
		1051		-19.7	1694	22.2
	Total	23	-54.7 -18.3	1722	20.6

Table D . 2

 D2 Sub-sample 09: Fluid inclusion data from microthermometry and Brillouin Spectroscopy with calculated salinity values afterBodnar (1993).Table D.3 Sub-sample 07: Fluid inclusion data from microthermometry andBrillouin Spectroscopy with calculated salinity values afterBodnar (1993).

	FI T eu T ice w biphasic FI T eu T ice w biphasic	Salinity Salinity
		[°C] [°C]	[°C] [°C]	[m/s] [m/s]	[wt. % NaCl eq.] [wt. % NaCl eq.]
	4 6	-51 -31.1 -25.3	1741 1876	29.4 25.8
	15 -54.7 -17.9 20 -54.2 -29.5	1680 1853	20.9 28.3
	18 -50 -22.9 31 -51.2 -32.7	1749 1837	24.3 30.4
	20 -52.7 -22.1 32 -55.2 -24	1692 1864	23.8 25.0
	21 -49 -19.6 44 -61 -31.7	1691 1810	22.1 29.8
	22 -69 -35.3 47 -33.5	1745 1804	32.3 31.0
	25 -58.5 -24.3 48 -31.5	1694	25.1 29.6
	30 52 -70 -27.3 _ -24	1699	25.0 27.0
	33 -53.1 -19.8 53 -57 -32	1690 1851	22.2 30.0
	35 -57 -27.4 61 -64 -30.8	1701 1917	27.1 29.2
	37 62 -73 -30.6 _ -25.6	1704 1885	25.9 29.0
	43 65 -66 -29.8 _ -12.6	1635 1865	16.5 28.5
	44 -51.4 -12.9 66 -31.6	1638 1862	16.8 29.7
	45 67 -72 -29.5 _ -20	1630 1851	22.4 28.3
	46 -54.7 -18.8 68 -61 -28.9	1685 1862	21.5 28.0
	48 -51 69	-26 -35.7	1707 1890	26.2 32.6
	49 -57 -28.3 71 -70 -30.5	1750 1815	27.6 29.0
	50 -57.7 -22.7 73 -29.1	1713 1825	24.1 28.1
	52 -54.2 -27 Total 17 -63.4 -30.2	1760 1848	26.8 28.8
	53	_	-33.9	1773	31.3
	Total 20 -54.7 -23.9	1707	24.7

T x mean + CI 95%

length of sampling interval in core section

mean

ACKNOWLEDGEMENTS

We thank M. Schwab and A. Brauer from GFZ Potsdam (Germany) and the Dead Sea Deep Drilling Project members for providing core material of sediment core DSDDP-5017-1-A. We appreciate the constructive feedbacks of two anonymous reviewers who largely improved the manuscript. This study is part of the PhD thesis of NB and was funded by Campus France during the national research program MOPGA.

Chapter 2

Sampling locations

Author Contributions NSB, EG, VG and FC developed the concept of the study. EG sampled the bulk core material. NSB sampled the core section and prepared the samples. NSB conducted the measurements and produced the entrapment temperature data. FC, EG and NSB analysed the data. NSB created the Ągures. All authors contributed to the main text, interpretations of data and Ąnalised the manuscript.

Competing Interests statement

The authors declare no competing interests.

Chapter 5. BIBLIOGRAPHY [START_REF] Wirth | A 2000 year long seasonal record of Ćoods in the southern European Alps[END_REF] Sharp increase of lake temperatures over one decade during interglacial MIS 7c: Insights from the Dead Sea basin Chapter 6. Sharp increase of lake temperatures over one decade during interglacial MIS 7c: Insights from the Dead Sea basin Typical pristine (non-altered) halite crystals contain normally growth-zonations with entrapped assemblages of FIs which appear macroscopically as visible cloudy bands (see Figure 1.1). Cumulate crystals contain also growth bands of FIs on a smaller scale with a FI-free central part that probably served as crystal nucleus.

FIs in CU layers are much smaller compared to CO layers, typically not exceeding sizes of 5 µm. Moreover, the FIs clouds are very dense in terms of the amount of FIs which complicates the temperature measurement with BS due to larger uncertaintes of the resulting T x data (see also Chapter 5).

Therefore, only coarse halite FIs have been studied for palaeothermometry. The selected FIs are located along dark-to-bright alternated growth bands (Figure 6.3 C-1). Nonetheless, it is important to note that not every single CO crystal has such growth zonations, so that some areas can be free of FIs. MIS 7c: Insights from the Dead Sea basin 6.4 Results

Entrapment temperatures

The total number of T x measurements per sample (n sample ) depends on the amount of trapped FIs that are in agreement with certain criteria (see Guillerm et al., 2020). Thus, n sample varies along the core section 176-2 because some selected sub-samples did not always contain suitable FIs in size, shape, or abundance.

A total of 13 coarse layers was studied, each one separated by Ąne cumulate layers with varying thicknesses. Here, the T x data are reported as mean values of each studied sub-sample, with uncertainties reported as standard error of the mean (SE).

The whole data set can be found in Table 6.1, results of both, CO-TOP sections are presented in Table 6.3 and of CO-BOT in Table 6.4, respectively. In addition, the results of CO-TOP sub-intervals have been split into two parts with respect to the core section, including data from both, 6-32 cm (upper part) and 39-63 cm (lower part), respectively (Table 6.6).

The coarse-TOP temperature is always very close to or higher than the BOT one, consistent with the expected seasonal warming (Table 6.1; Figure 6.4). There is an excursion in layer 8 with an anomalous increase in coarse-BOT temperature;

however the two values of layer 8 have marginally overlapping error bands, which leaves the possibility for the two values to be nearly equal. The coarse-TOP temperatures are Ąrst nearly constant and then increase, while the BOT temperatures show a 3-year cycling of relatively lowest temperatures (Figure 6.5).

Exclusively all sub-samples stratigraphically above CO8-TOP show a rise in T x along the growth direction of coarse layers (Figure 6.4 A), further, all remaining CO-TOP data are signiĄcantly elevated compared to sub-samples from CO13-TOP to CO8-TOP with a mean of 19.0 °C vs. 17.0 °C for the uppermost 26 cm and the lowermost 24.2 cm, respectively (Tables 6 .5, 6.6). Hence, the interval from CO13-BOT to CO8-TOP is deĄned as lowermost interval I low (445.695 to 445.423 m blf; 27.3 cm length).

Chapter 6. Sharp increase of lake temperatures over one decade during interglacial MIS 7c: Insights from the Dead Sea basin

Halite layer thicknesses

The absolute thickness ∆d abs of halite layers differs from that of the calculated sub-sample locations (CO-TOP minus CO-BOT) because the samples have not been taken exactly at the grain size boundaries but rather a few millimeters above.

Therefore, the following halite thicknesses are reported as ∆d abs .

The mean thickness of all coarse layers is 2.6 cm. The youngest coarse layer CO1

has the maximum thickness (3.5 cm), while the minimum thickness was measured in the second-youngest coarse layer CO2 (1.7 cm). The three deĄned intervals have mean thicknesses of 2.7 cm (I low ), 2.5 cm (I cen ), and 2.4 cm (I upp ), respectively.

Cumulate halite layers have a mean thickness of 2.1 cm with a minimum value of 0.6 cm between CO6-TOP and CO5-BOT, and a maximum thickness of 4.3 cm between CO8-TOP and CO7-BOT. The minimum cumulate layer is located within the sub-interval where both CO-BOT and CO-TOP data show a Lmin.

Furthermore, note that the maximum cumulate thickness coincides with the sharp increase in T x from CO7-BOT to CO7-TOP. Finally, each of the deĄned intervals yields a mean cumulate layer thickness of 2.4 cm (I low ), 2.2 cm (I cen ), and 1.6 cm (I upp ), respectively.

Chapter 6. Sharp increase of lake temperatures over one decade during interglacial MIS 7c: Insights from the Dead Sea basin Figure 6.9 Modelled changes of mean summer maximum (July-August, JJA; TX) and mean winter minimum (December-February, DJF; TN) temperatures. Changes relative to a control period for 2040Ű2069 (left) and 2070Ű2099 (right) (Lelieveld et al., 2012). Tzedakis et al. (2004) showed positive isotopic excursions in both benthic and planktonic foraminifera from a marine core off SW Portugal. Unfortunately, terrestrial climate archives are scarce in the Levant region. Speleothems from the Soreq Cave have shown to be able to reconstruct the climatic history of the last 160 ka, covering MIS 6 (Matthews et al., 2021), and to decipher modern seasonal changes on monthly time scales (Orland et al., 2014). Therefore, the results obtained here from a halite section covering 13 years of lake history might encourage further palaeotemperature reconstructions by using Brillouin spectroscopy. To uncover additional seasonal patterns of other interglacial periods will improve our understanding of the Dead Sea Basin, and extend the database derived from palaeoclimate proxies.

Chapter 6. Sharp increase of lake temperatures over one decade during interglacial MIS 7c: Insights from the Dead Sea basin

The differences between T x of both, CO-BOT(n+1) and CO-BOT(n) (pre) could be used to describe the evolution of annual temperature deĄcits prior to lake warming during spring-summer along the core section. On the opposite, the difference between T x of both, CO-TOP(n+1) and CO-TOP(n) (post) may be used to identify lake temperature deĄcits during the post-summer period. The lake cooling during winter may be inferred from the difference between CO-TOP(n)

and CO-BOT(n+1), and a potential lake warming upon summer season from CO-TOP(n) and CO-BOT(n) (and for n+1, etc.) (Figure 6.10 C).

Although other proxy data are lacking in the Levantine region for MIS 7, the mean of each T x of CO-TOP(n, n+1) and CO-BOT(n, n+1) might be useful to compare along the core section. Therefore, the mean annual temperature writes:

The mean winter temperature is obtained by calculating the mean of CO-TOP(n)

and CO-BOT(n+1), respectively.

Chapter 7

Entrapment temperatures of gypsum To summarise, several cooling-heating cycles are often needed to trigger a transformation from liquid to liquid+vapor in natural gypsum samples. The recommended procedure is to cool the sample at intermediate rates (e.g., 10 °C/min) to -60 °C for ca 2-3 minutes, followed by heating to ca 30 °C at the same rate. If needed, the cycle might be repeated until the remaining FIs cavitated, which had been previously measured in monophasic state and further selected for BS analysis.

Chapter 7. BIBLIOGRAPHY Sorrel, P., Popescu, S.-M., Head, M., Suc, J., Klotz, S., and Oberhänsli, H. (2006). (1) <1600 m/s for low-saline samples,

Hydrographic development of the

(2) <1700 m/s for intermediary saline samples, and

(3) >1700 m/s for high-saline samples (Figure 8.2). A selenite-type gypsum samples has been collected from a saline sabkha, Tunisia, and further split into three sub-samples (labelled as "05", "07", and "09", respectively). The studied host crystal shows a typical growth direction with a V-shaped structure, consisting of cloudy, clear, and intermediate intervals, as it had been reported for Messinian twinned selenite (Figure 8.7;Natalicchio et al., 2014;Costanzo et al., 2019). Following this, the lower sub-sample "07" marks the relatively oldest part with respect to sub-samples "09" (central part) and "05" (upper/youngest part).

This section of the thesis aims to compare sound velocity data sets taken from these three intervals, and to test if the sound velocity varies signiĄcantly along the growth direction of the selenite host crystal. Following one of the classical approaches of evaporating brines (e.g., Warren, 2016), the hypothesis is that a brine becomes more saline the more water evaporates, leading to more saline entrapped Ćuid inclusions in a crystal over a longer temporal distance. On the contrary, a brine could have been diluted over time by rainfall, groundwater inĆow, or marine Ćooding, which would lower the brineŠs salinity. Therefore, this study is limited to the Ąrst application of the BS method to gypsum only, in order to provide the fundamental knowledge for further projects. Since the growth duration of the studied sample is unknown and other methods could not have been applied (e.g., δ 34 S, δ 18 O, δD isotopic composition; elemental concentrations), the interpretation of the obtained data set remains tentative.

Results

A total of 284 Ćuid inclusions were measured at constantly 25 °C, and Ąnal data range from 1468 to 1975 m/s. By grouping the data set according to each subsample, a strong contrast is observed in the statistics for mean, median, standard deviation (SD), standard error (SE), minimum (min), and maximum (max) values (Table 8.3). Sub-sample "05" has a mean of 1644 ±79 (SD) m/s (n = 185) versus 1837 ±54 m/s (SD) of sub-sample "07" (N = 98). The minimum value of "05" ("07") is 1468 m/s (1739 m/s), the maximum is about 1877 m/s (1975 m/s).

Further ordering by both FI type (primary, secondary, mono-or biphasic) and thermodynamic phase (primary biphasic) of each grouped data set shows mean offsets of 184 m/s for primary monophasic (n = 107 and 49 for "05" and "07", respectively), 227 m/s for secondary monophasic (n = 39 and 13 for "05" and "07", respectively), and 141 m/s for primary biphasic Ćuid inclusions (n = 39 Figure 8.8 H, the data distribution is broader in sample "05" which is observed also in selected FIs (Figure 8.9). The mean of sample "09" is slightly lower than that of "05" and "07" while its range is the lowest of all sub-samples (Table 8.4).

Moreover, a step-wise decrease in minimum sound velocity values coincides with the growth direction of the host gypsum sample with sub-sample "05" being the youngest part, "09" the intermediate part, and "07" the oldest part with respect to the location inside the whole sample. The minimum value of "05" (1550 m/s) equals the calculated capillary value of 5 wt.% NaCl, whereas "09" (1630 m/s) is close to the capillary value of >10 wt.% NaCl, and "07" (1754 m/s) already well above 20 wt.% NaCl (see Table 8.2). Crisis -a period of massive deposition of evaporites (mainly salt and gypsum) in the Mediterranean Sea -two major studies have suggested that selenite gypsum types might have rather formed in low-saline, freshwater-dominated environments (Natalicchio et al., 2014), or in usual marine environments (Bigi et al., 2022).

The latter study revealed that FIs might be likely affected by a crack-and-seal mechanism which could directly alter the composition of the trapped brine.

Further, the authors hypothesized that this process could affect as well the geochemistry of the crystallization water, making gypsum-rich FIs not very suitable for palaeosalinity studies (Bigi et al., 2022). However, the same authors also proposed an extended temperature excursion (220 °C from -100 to +120 °C) In this preliminary study, only the determination of Ąrst melting temperatures ("eutectic temperatures"; T e ) of FIs can help to classify the aqueous system (Bodnar, 1993). Here, nearly all observed eutectic temperatures fall below that one of the NaCl Ű H 2 O system (-21.2 °C). However, as described in Goldstein and Reynolds (1994), the identiĄcation of the true system ̸ = NaCl is often ambitious and problematic if daughter minerals appear.

Therefore, the obtained T ice data were used only for simpliĄcation because the aim of this work is not to reconstruct the absolute ("true") salinity of single FIs, but rather to test if relative salinity variations occur along the growth direction of the gypsum host crystal. By comparing the data set with other experimental data obtained during the PhD thesis, the apparent maximum sound velocity data are observed in sub-sample "07" (oldest part). On the contrary, lowest sound velocity data are found in ( 1) two data sets derived from secondary FIs from Messinian sampling locations, which reĆect likely inĄltrating freshwater through time, and

(2) in a lab-grown gypsum sample (pure water FIs) (Figure 8.12). Therefore, it can be argued that sub-sample "07" originated in a highly saline brine that was later slightly diluted but still above the gypsum saturation point. This is underlined by both sub-samples "09" and "05" which yield sound velocity values below the NaCl saturation line (red dotted curve in Figure 8.12). The Ąnal comparison with the theoretical "pure" NaCl curve shows a deviation in both salinity and sound velocity values, which might be explained by the non-binary system of the Ćuid inclusions of the natural gypsum sample. The results need to be re-evaluated by future studies including additional methods to account for the brine composition. Furthermore, a larger data set should be produced by measuring ≥ 50 biphasic FIs in each crystal, being aligned along the same growth bands. In addition, the more complex brine composition should be modeled to compare both, the measured and the theoretical salinity of gypsum FIs.

Nevertheless, the results give hope to use gypsum from outcrops, bore cores, or arid sabkha environments, to estimate the palaeosalinity of a well-dated sample.

It is recommended to setup a bunch of laboratory-controlled standard samples at varying salinity, which then could be used to compare natural samples after analysis with Brillouin spectroscopy. The advantage of this method is the fast acquisition of large data sets, since only biphasic FIs need to be measured at one constant temperature (e.g., 25 °C). This reduces the lab costs because it avoids the usage of liquid nitrogen resources. 4. Switch on both, the screen and the camera.

5. Brillouin box: press both "power on" buttons in front of the interferometer controls right above the black Brillouin box. At the backside, turn the "detector" bottom to ON. Tipp: place all three software windows to the left screen in order to have an overview while working with the microscope! Make sure that you have put the same channel number at the interferometer control as set in "edit comments", and be aware that the Ąlter conĄguration for the reference light beam (behind the black Brillouin box) has likely to be changed in order to get an appropriate baseline.

Appendix B

Ruling out thermal effects due to the laser beam

B.1 Method

One way to speed up measurements is to increase the laser intensity on the sample.

In order to check that this is not the case, experiments were performed to compare the values of speed of sound for selected FIs in synthetical NaCl samples with known entrapment temperatures.

with IF new the desired laser intensity on the sample in mW, S A the scan amplitude (for 37.57 GHz = 0.3757) and IF m the laser intensity to be measured by using a power meter. Entrapment temperatures from natural gypsum samples