Changle Wang
email: wangcl@mail.iggcas.ac.cn

Leslie J Robbins
email: leslie.robbins@uregina.ca

Noah J Planavsky

Nicolas J Beukes

Laureline A Patry

Stefan V Lalonde

Maxwell A Lechte

Dan Asael

Christopher T Reinhard

Lianchang Zhang

Kurt O Konhauser

Archean to early Paleoproterozoic iron formations document a transition in iron oxidation mechanisms

teaching and research institutions in France or abroad, or from public or private research centers.

Introduction

The atmosphere during the first half of Earth's 4.5-billion-year (Ga) history was essentially devoid of free oxygen (O2), with atmospheric abundances less than 0.001% of the present atmospheric level (PAL) [START_REF] Farquhar | Geological constraints on the origin of oxygenic photosynthesis[END_REF][START_REF] Lyons | The rise of oxygen in Earth's early ocean and atmosphere[END_REF] .

Geological and geochemical constraints reveal that the first permanent rise in the partial pressure of atmospheric O2 (pO2) occurred between ~2.5 and 2.3 Ga, a period that has been termed the Great Oxidation Event (GOE) [START_REF] Bekker | Dating the rise of atmospheric oxygen[END_REF][START_REF] Canfield | The early history of atmospheric oxygen: homage to Robert M. Garrels[END_REF][START_REF] Guo | Reconstructing Earth's surface oxidation across the Archean-Proterozoic transition[END_REF][START_REF] Holland | The oxygenation of the atmosphere and oceans[END_REF][START_REF] Luo | Rapid oxygenation of Earth's atmosphere 2.33 billion years ago[END_REF][START_REF] Warke | The Great Oxidation Event preceded a Paleoproterozoic "snowball Earth[END_REF] . Despite widespread agreement on these general trends, the tempo and mode of Earth's early oxygenation, and the redox conditions leading up to the GOE, remain poorly understood as few redox proxies are sufficiently precise to track secular variations in ocean-atmosphere O2 abundance. Although incipient oxygenation of Earth's surface environments before GOE has commonly been proposed [START_REF] Anbar | A whiff of oxygen before the Great Oxidation Event?[END_REF][START_REF] Koehler | Transient surface ocean oxygenation recorded in the∼ 2.66-Ga Jeerinah Formation, Australia[END_REF][START_REF] Planavsky | Evidence for oxygenic photosynthesis half a billion years before the Great Oxidation Event[END_REF] , there still remains an active debate regarding the exact timing of the first emergence of oxygenic photosynthesis [START_REF] Rasmussen | Evidence for anoxic shallow oceans at 2.45 Ga: Implications for the rise of oxygenic photosynthesis[END_REF][START_REF] Soo | The evolution of photosynthesis and aerobic respiration in the cyanobacteria[END_REF] , with suggestions that this biological innovation may have coincided with the oxygenation of the atmosphere [START_REF] Rasmussen | Evidence for anoxic shallow oceans at 2.45 Ga: Implications for the rise of oxygenic photosynthesis[END_REF][START_REF] Fischer | Evolution of oxygenic photosynthesis[END_REF] . Although, a recent phylogenomic study on enzymes that either produce or utilize oxygen points to an Archean origin for oxygenic photosynthesis [START_REF] Jabłońska | The evolution of oxygen-utilizing enzymes suggests early biosphere oxygenation[END_REF] . In much the same manner, various geochemical proxies likely have different sensitivities to changes in pO2 and may not respond uniformly [START_REF] Anbar | A whiff of oxygen before the Great Oxidation Event?[END_REF][START_REF] Koehler | Transient surface ocean oxygenation recorded in the∼ 2.66-Ga Jeerinah Formation, Australia[END_REF] , further blurring the history of atmospheric and oceanic pO2 as well as the exact timing of the GOE. This obscurity in the onset and timing of Earth's early oxidation has also given rise to recent questions as to what role biological oxidation played, if any, in the deposition of iron formations (IFs).

Iron formations are particularly useful sedimentary archives for tracking Fe sources and sinks during much of the Archean and early Paleoproterozoic as they formed in a range of depositional settings on the continental shelves of existing cratons [START_REF] Bekker | Iron Formation: The Sedimentary Product of a Complex Interplay among Mantle, Tectonic, Oceanic, and Biospheric Processes[END_REF][START_REF] Konhauser | Iron formations: A global record of Neoarchaean to Palaeoproterozoic environmental history[END_REF] . Moreover, because IF are chemical sediments, they represent an archive of environmental conditions and seawater composition at the time of their deposition [START_REF] Konhauser | Iron formations: A global record of Neoarchaean to Palaeoproterozoic environmental history[END_REF] , which in turn, provides insights into the availability of trace-element nutrients to, and the activity of, the ancient biosphere [START_REF] Robbins | Trace elements at the intersection of marine biological and geochemical evolution[END_REF] .

The best-preserved IF successions have a remarkably uniform mineralogy, composed mostly of chert, magnetite, and hematite, with variable amounts of Fe-rich silicate minerals (e.g., greenalite), carbonate minerals (e.g., siderite), and locally sparse sulfides (e.g., pyrite). It is also apparent that none of the minerals in IFs today are primary in origin, and that the original seafloor precipitate mineralogy was not preserved. Instead, the observed minerals reflect multiple post-depositional alteration processes that modified the primary precipitates which are generally believed to have been ferric oxyhydroxides (e.g., ferrihydrite) and amorphous silica [START_REF] Konhauser | Iron formations: A global record of Neoarchaean to Palaeoproterozoic environmental history[END_REF] (see Supplementary Materials for a detailed discussion on the primary mineralogy of IF).

The presence of ferric iron (Fe(III)) minerals in IFs is generally ascribed to the metabolic activity of planktonic bacteria in the photic zone of Earth's ancient oceans. The classic model invokes Fe(III) precipitation occurring at the interface between oxygenated shallow waters and reduced upwelling ferrous iron (Fe(II))-rich waters, the oxygen (O2) being sourced from cyanobacteria, or their predecessors [START_REF] Cloud | Paleoecological significance of the banded iron-formation[END_REF] . These photoautotrophs would have flourished when nutrients were available and passively induced the precipitation of Fe(III)-oxyhydroxide through the excretion of their metabolic waste (i.e. O2). Alternatively, anoxygenic photoautotrophic Fe(II)-oxidizing bacteria (known as photoferrotrophs) could have directly oxidized Fe(II) and coupled this to the fixation of carbon utilizing light energy [START_REF] Garrels | Genesis of Precambrian iron-formations and the development of atmospheric oxygen[END_REF] . From a mass balance perspective, either metabolism could account for all the Fe(III) deposited in BIF [START_REF] Konhauser | Could bacteria have formed the Precambrian banded iron formations?[END_REF][START_REF] Konhauser | Phytoplankton contributions to the trace-element composition of Precambrian banded iron formations[END_REF] . The classical view, however, has recently been challenged, with the suggestion that IFs were largely deposited as Fe(II) silicates (see [START_REF] Muhling | Widespread deposition of greenalite to form Banded Iron Formations before the Great Oxidation Event[END_REF] ), a model which would negate the need for oxygenic or anoxygenic photosynthesis in IF deposition. Inspired by the apparent necessity for biologically driven redox reactions and IF precipitation, and its potential to provide insights into the oxygenation of Earth in the lead-up to the GOE, here we explore the coupled proxies of Fe isotope and manganese (Mn) to Fe ratios in an effort to better understand the link between IF deposition and the rise of oxygenic photosynthesis during the Archean and Paleoproterozoic.

Iron Formation Deposition and Marine Oxygen Levels

As Fe in seawater is soluble under anoxic conditions as Fe(II) and is rendered poorly soluble under oxic conditions, precipitating readily as Fe(III) oxyhydroxides, the secular record of Fe oxidation may potentially be leveraged to track surface O2 levels [START_REF] Canfield | The early history of atmospheric oxygen: homage to Robert M. Garrels[END_REF][START_REF] Czaja | Biological Fe oxidation controlled deposition of banded iron formation in the ca. 3770 Ma Isua Supracrustal Belt (West Greenland)[END_REF] . One potential means by which we can parse the mechanisms underpinning Fe(II) oxidation in the ancient oceans is through measurement of the relative abundances of the different stable isotopes of Fe in the Fe-oxides of IFs 24 . This approach is appealing since the kinetics of Fe(II) oxidation are well-constrained and the Fe isotope systematics of Fe(II) oxidation by photoferrotrophs and ambient O2 are relatively well-understood [START_REF] Busigny | Iron isotopes in an Archean ocean analogue[END_REF][START_REF] Dauphas | Iron isotope systematics[END_REF][START_REF] Johnson | The Ancient Earth[END_REF][START_REF] Croal | Iron isotope fractionation by Fe (II)oxidizing photoautotrophic bacteria[END_REF] . Briefly, partial oxidation of dissolved Fe(II) produces Fe(III)-oxyhydroxides that are enriched in [START_REF] Rouxel | Iron isotope constraints on the Archean and Paleoproterozoic ocean redox state[END_REF] Fe, typically by 1 to 3‰, leaving the residual dissolved Fe(II) [START_REF] Rouxel | Iron isotope constraints on the Archean and Paleoproterozoic ocean redox state[END_REF] Fe-depleted, although more rapid oxidation effectively lowers this overall fractionation [START_REF] Johnson | The Ancient Earth[END_REF] . This contrasts with the formation of reduced iron phases (e.g., Fe carbonates and silicates): these pathways are typically considered to involve either a negligible fractionation or a small negative fractionation [START_REF] Dauphas | Iron isotope systematics[END_REF][START_REF] Johnson | The Ancient Earth[END_REF] , leaving residual dissolved Fe(II) slightly [START_REF] Rouxel | Iron isotope constraints on the Archean and Paleoproterozoic ocean redox state[END_REF] Fe-enriched. Furthermore, when Fe(II) oxidation is complete and quantitative-where there is a sharp redoxcline and a well-oxygenated surface layer-there is no net fractionation during the formation of Fe(III) oxides [START_REF] Busigny | Iron isotopes in an Archean ocean analogue[END_REF] . However, the fractionation factors of Fe isotopes for Fe(II) oxidation by free O2 and anoxygenic photosynthesis effectively overlap 29 , making it difficult to clearly recognize the mechanism responsible for Fe(II) oxidation at different points in Earth's history based on solely Fe isotopes. To address this shortcoming, a supplementary record that can be directly tied to Fe(II) oxidation is required. Similar to Fe, manganese (Mn) is redox-sensitive and soluble under anoxic conditions. The oxidation of Mn(II) and the subsequent formation of Mn(III/IV) oxides (from herein simply referred to as Mn oxides) in the presence of free O2 generally requires both higher redox potentials than the oxidation of Fe(II) and the catalyzing activity of a Mn(II)-oxidizing bacteria [START_REF] Tebo | Geomicrobiology of manganese (II) oxidation[END_REF] . Manganese oxides would also be expected to undergo rapid reductive dissolution in anoxic, Fe(II)-rich seawater. As such, Mn enrichments resulting from either Mn oxide precipitation or a Mn-oxide shuttle to the sediment-seawater interface may be linked to the presence of an oxygenated water column [START_REF] Planavsky | Evidence for oxygenic photosynthesis half a billion years before the Great Oxidation Event[END_REF][START_REF] Kurzweil | Manganese oxide shuttling in pre-GOE oceans-evidence from molybdenum and iron isotopes[END_REF] . Although the sedimentary and diagenetic pathways responsible for the incorporation of Mn into IFs has recently been questioned [START_REF] Siahi | Insights into the processes and controls on the absolute abundance and distribution of manganese in Precambrian iron formations[END_REF] , secular trends in the relative enrichment of Mn in IFs are typically tied to marine redox conditions. Therefore, as IFs serve as robust archives of both the Fe isotope and Mn/Fe ratios in the deep time, they may be interrogated to constrain the redox state of local seawater during IF deposition. Importantly, a coupled approach such as this can provide new insights into the presence of free O2 in seawater which, in turn, argues against possible anaerobic pathways capable of oxidizing Mn(II) [START_REF] Daye | Light-driven anaerobic microbial oxidation of manganese[END_REF][START_REF] Liu | Anoxic photogeochemical oxidation of manganese carbonate yields manganese oxide[END_REF] . Such insights are strengthened when a correlation between these two parameters exists. That is to say, these coupled records are capable of tracking the timing and the evolution, of O2-driven Fe(II) and Mn(II) oxidation across Earth's history.

Tracking Iron Oxidation with Iron Isotopes

Here, we present a detailed record of bulk-rock Fe isotopes (reported as the ratio of 56 Fe/ [START_REF] Reinhard | Oxidative sulfide dissolution on the early Earth[END_REF] Fe, relative to the IRMM-014 reference material standard in parts per thousand; δ 56 Fe) and Mn/Fe ratios from Archean to early Paleoproterozoic IFs in order to explore the redox dynamics of the early Earth. Data derived from ~770 individual samples include both new results (n = 253) and those obtained from a comprehensive literature survey (n = 517). Samples were passed through a series of strict filters to select for IFs that are most likely to retain a seawater signature (see Supplementary Materials). Nearly all samples have a negligible detrital component (< 0.5 wt.% Al2O3) and there exists no significant correlation between the paleoredox data (δ [START_REF] Rouxel | Iron isotope constraints on the Archean and Paleoproterozoic ocean redox state[END_REF] Fe values and Mn/Fe ratios) and detrital indicators (e.g., Al2O3, Table S3), suggesting that secular variations in redox indicators were not driven by detrital input. The predominance of positively fractionated Fe isotope values in IFs is supportive of genetic models, which invoke precipitation of primary Fe(III)-oxyhydroxide minerals [START_REF] Dauphas | Iron isotope systematics[END_REF] , as opposed to primary Fe-silicate minerals [START_REF] Rasmussen | Evidence for anoxic shallow oceans at 2.45 Ga: Implications for the rise of oxygenic photosynthesis[END_REF][START_REF] Muhling | Widespread deposition of greenalite to form Banded Iron Formations before the Great Oxidation Event[END_REF] . Indeed, a recent evaluation of triple Fe isotope systematics in Neoarchean to Paleoproterozoic IF and pyrite samples, indicates that the Neoarchean would have been characterized by a significant oxidized Fe sink prior to the GOE [START_REF] Heard | Triple iron isotope constraints on the role of ocean iron sinks in early atmospheric oxygenation[END_REF] . Aside from being consistent with the Fe isotopic record, the requirement of extensive Fe oxidation is supported by a recent model that has challenged whether post-depositional groundwater flow could have oxidized a significant volume of Fe(II) on reasonable time scales [START_REF] Robbins | Hydrogeological constraints on the formation of Palaeoproterozoic banded iron formations[END_REF] . Therefore, we suggest that IF Fe isotope systematics are consistent with the interpretation that the marine biosphere, namely Fe-oxidizing bacteria (e.g., photoferrotrophs or cyanobacteria), were involved in the oxidation of Fe(II) and the subsequent deposition of IFs. Importantly, Fe isotope data also suggest that secular changes may present a valuable tool for tracking changes in the relative importance of the mechanisms underlying marine Fe(II) oxidation.

To facilitate the discussion of our data, we have split the deposition of IFs into five stages based upon secular variations in their δ [START_REF] Rouxel | Iron isotope constraints on the Archean and Paleoproterozoic ocean redox state[END_REF] Fe values and Mn/Fe ratios (Fig. 1). Although these bins did not emerge from a robust statistical treatment of the data, this binning helps illustrate that there are fundamental shifts in the redox evolution of Earth's early seawater and atmosphere. Stage 1 extends from the Eoarchean into the Paleoarchean (>3.3 Ga) and is characterized by highly positive δ [START_REF] Rouxel | Iron isotope constraints on the Archean and Paleoproterozoic ocean redox state[END_REF] Fe values (Fig. 1a), likely the product of deposition under oxygen-limited conditions that favored the partial oxidation of a relatively large marine reservoir of aqueous Fe(II). Also characterized by low Mn/Fe ratios (<0.01) (Fig. 1b), Stage 1 IFs reflect deposition in an anoxic ocean-atmosphere system. This is consistent with evidence from the sedimentary sulfur isotope record, namely the mass-independent fractionation of sulfur isotopes (MIF-S), which is the prevailing geochemical index for an anoxic atmosphere [START_REF] Farquhar | Geological constraints on the origin of oxygenic photosynthesis[END_REF][START_REF] Farquhar | Atmospheric influence of Earth's earliest sulfur cycle[END_REF] [START_REF] Rouxel | Iron isotope constraints on the Archean and Paleoproterozoic ocean redox state[END_REF] Fe values (Fig. 1a), suggesting either a greater degree of anoxygenic photosynthetic Fe(II) oxidation [START_REF] Czaja | A multistage origin for Neoarchean layered hematite-magnetite iron formation from the Weld Range, Yilgarn Craton, Western Australia[END_REF] or elevated marine O2 concentrations. In conjunction with higher Mn/Fe ratios in Stage 2 IFs (Fig. 1b), we propose that the transition to Stage 2 coincides with the emergence and/or increased impact of oxygenic photosynthesis, implying that free O2-driven Fe(II) oxidation has been active since ~3.3 Ga. Stage 2 encompasses the deposition of four successions suggested to preserve evidence of pre-GOE O2 oases (i.e. oxygenated shallow seawater)-the 3.2 Ga Fig Tree and Moodies groups in the Barberton Greenstone Belt 39, 40 , the 2.95 Ga Sinqeni Formation in the Pongola Supergroup [START_REF] Planavsky | Evidence for oxygenic photosynthesis half a billion years before the Great Oxidation Event[END_REF] , and the 2.94 Ga Red Lake Greenstone Belt. The data from Moodies and Sinqeni IFs also feature a significant negative correlation between Mn/Fe ratios and δ [START_REF] Rouxel | Iron isotope constraints on the Archean and Paleoproterozoic ocean redox state[END_REF] Fe values (Fig. S1), supporting the interpretation that free O2 began to accumulate in shallow-marine settings during this interval, allowing for Mn(II) oxidation. In the presence of such a shallow-water oxic zone, photoferrotrophs would still have thrived, growing at depth near the base of the photic zone [START_REF] Kappler | Deposition of banded iron formations by anoxygenic phototrophic Fe (II)-oxidizing bacteria[END_REF] . This would allow photoferrotrophs a preferential access to upwelling nutrients owing to their adaption to lower light levels, thus limiting proliferation of cyanobacteria in the surface ocean [START_REF] Jones | Iron oxides, divalent cations, silica, and the early earth phosphorus crisis[END_REF] and confining cyanobacteria to very near-shore environments where there may have been nutrient input from emergent continents. Further, cyanobacteria expansion would have been held in check by the high Fe(II) concentrations in the upwelling currents because of the oxidative stress resulting from reactions between Fe(II) and O2 [START_REF] Swanner | Modulation of oxygen production in Archaean oceans by episodes of Fe (II) toxicity[END_REF] . Indeed, marginalization of cyanobacteria to shallow waters is consistent with limited O2 production in the Mesoarchean ocean, as evidenced by nitrogen (N) isotopes from shales in the Pongola Supergroup [START_REF] Ossa | Limited oxygen production in the Mesoarchean ocean[END_REF] and the lack of a correlation between δ 56 Fe values (up to 1.5‰) and Mn/Fe ratios within the Red Lake banded IF (BIF) (Fig. S1), implying that the redoxcline remained likely above the photic zone. Therefore, the contribution of microaerophilic and nitrate-reducing Fe(II)-oxidizers to Fe cycling [START_REF] Melton | The interplay of microbially mediated and abiotic reactions in the biogeochemical Fe cycle[END_REF] during this stage is considered to be minor. Moreover, while a sharp redox boundary between deep and shallow waters was proposed to exist at the depositional site of IFs from the 3.2 Ga Fig Tree Group [START_REF] Satkoski | A redox-stratified ocean 3.2 billion years ago[END_REF] , subsequent studies have reported the frequent occurrence of highly positive δ [START_REF] Rouxel | Iron isotope constraints on the Archean and Paleoproterozoic ocean redox state[END_REF] Fe values for these IFs [START_REF] Busigny | Iron and sulfur isotope constraints on redox conditions associated with the 3.2 Ga barite deposits of the Mapepe Formation (Barberton Greenstone Belt, South Africa)[END_REF][START_REF] Planavsky | Iron isotope composition of some Archean and Proterozoic iron formations[END_REF] , likely supporting either partial Fe(II) oxidation in a reducing oceanic environment [START_REF] Busigny | Iron and sulfur isotope constraints on redox conditions associated with the 3.2 Ga barite deposits of the Mapepe Formation (Barberton Greenstone Belt, South Africa)[END_REF][START_REF] Planavsky | Iron isotope composition of some Archean and Proterozoic iron formations[END_REF] or heterogeneity in ocean redox conditions [START_REF] Johnson | The Ancient Earth[END_REF] . It is noted that the ~2.95 Ga granular IF (GIF) in the upper Nconga Formation of the Pongola Supergroup possesses positive δ [START_REF] Rouxel | Iron isotope constraints on the Archean and Paleoproterozoic ocean redox state[END_REF] Fe values and low Mn/Fe ratios [START_REF] Smith | Oncoidal granular iron formation in the Mesoarchaean Pongola Supergroup, southern Africa: Textural and geochemical evidence for biological activity during iron deposition[END_REF] . These features are in stark contrast with those observed in the O2 oases described above, implying that anoxic conditions were prevalent, disrupted by the presence of spatially restricted disequilibrium between the surface oxygenated oceans and reducing atmosphere [START_REF] Eickmann | Isotopic evidence for oxygenated Mesoarchaean shallow oceans[END_REF] . To this end, the 3.22 Ga Moodies Group also records depleted C and enriched N isotopes that suggest an oxidative biogeochemical cycle in a terrestrial environmental [START_REF] Homann | Microbial life and biogeochemical cycling on land 3,220 million years ago[END_REF] , consistent with O2 having been present in local environments near shallow marine oases.

The Fe isotopes and Mn/Fe ratios of IFs deposited during Stage 3 (~2.8-2.5 Ga) are comparable with those of Stage 1 (Fig. 1). Transient and spatially restricted oxygenation events have also been suggested to occur during this interval based on a wide range of geochemical proxies [START_REF] Koehler | Transient surface ocean oxygenation recorded in the∼ 2.66-Ga Jeerinah Formation, Australia[END_REF][START_REF] Czaja | Evidence for free oxygen in the Neoarchean ocean based on coupled iron-molybdenum isotope fractionation[END_REF][START_REF] Wille | Evidence for a gradual rise of oxygen between 2.6 and 2.5 Ga from Mo isotopes and Re-PGE signatures in shales[END_REF] . It also needs to be noted that sulfur-rich exhalites in the ~2.7 Ga Abitibi greenstone belt possess highly negative Fe isotope and higher Mn contents, suggesting the presence of O2 in the upper part of the water column [START_REF] Hiebert | Depositional setting of the Late Archean Fe oxide-and sulfide-bearing chert and graphitic argillite in the Shaw Dome, Abitibi Greenstone Belt, Canada[END_REF] . However, the IF data presented here do not record evidence of oxygenation events, likely due to either sampling bias or the relatively deep-water depositional environments (below storm wave base) of these IFs. Nevertheless, these samples are from fifteen separate, geographically widespread successions (see Supplementary Materials), and are thus considered to be fairly representative of this time interval. The common occurrence of positive δ [START_REF] Rouxel | Iron isotope constraints on the Archean and Paleoproterozoic ocean redox state[END_REF] Fe values indicates partial Fe(II) oxidation, coupled with lower Mn/Fe ratios, indicating the widespread presence of fully anoxic conditions. These conditions likely existed in the upper water column with photoferrotrophs being the dominant force in driving Fe(II) oxidation. There is no clear relationship between δ [START_REF] Rouxel | Iron isotope constraints on the Archean and Paleoproterozoic ocean redox state[END_REF] Fe values and Mn/Fe ratios during this stage (Fig. S1), further confirming the absence of free O2 in the water column. A lack of discrete oxygen oases during this period-as suggested by our data-is further consistent with an Archean world characterized by low levels of primary productivity [START_REF] Olson | Quantifying the areal extent and dissolved oxygen concentrations of Archean oxygen oases[END_REF][START_REF] Reinhard | Oxidative sulfide dissolution on the early Earth[END_REF] . Stage 4, from 2.5 to 2.4 Ga, is represented by the most laterally extensive deposition of IFs in Earth history [START_REF] Konhauser | Iron formations: A global record of Neoarchaean to Palaeoproterozoic environmental history[END_REF] , and is immediately preceded by a "whiff" of oxygen at 2.5 Ga 9, 55 . The most striking feature in this interval is that δ [START_REF] Rouxel | Iron isotope constraints on the Archean and Paleoproterozoic ocean redox state[END_REF] Fe values and Mn/Fe ratios have substantially decreased and increased, respectively, marking the gradual transition of the dominant IF oxidant from photoferrotrophs to free O2 (Fig. 1). Specifically, for the coeval and correlative Transvaal (South Africa) and Hamersley (Western Australia) basins, continuous IF records and precise geochronological constraints reveal a clear trend for these proxies (Fig. S9 andS12). This trend reflects changing marine redox conditions across large portions of the continental margins worldwide. The δ 56 Fe values in IFs become more negative towards the latter stage of Stage 4, while Mn/Fe ratios become higher. Significantly, negative correlations between Mn/Fe ratios and δ 56 Fe values are observed for all IF samples (Fig. S2). We suggest that this is the product of gradually increasing O2 levels within seawater, whereby the negative δ [START_REF] Rouxel | Iron isotope constraints on the Archean and Paleoproterozoic ocean redox state[END_REF] Fe values are linked to a progressive distillation during the drawdown of the Fe(II) reservoir [START_REF] Kurzweil | Manganese oxide shuttling in pre-GOE oceans-evidence from molybdenum and iron isotopes[END_REF][START_REF] Rouxel | Iron isotope constraints on the Archean and Paleoproterozoic ocean redox state[END_REF] , resulting in a [START_REF] Rouxel | Iron isotope constraints on the Archean and Paleoproterozoic ocean redox state[END_REF] Fe-depleted seawater reservoir. A microbially driven benthic Fe shuttle [START_REF] Severmann | Modern iron isotope perspective on the benthic iron shuttle and the redox evolution of ancient oceans[END_REF] -which preferentially releases light Fe isotopes-may have also contributed to this isotopically depleted Fe(II) reservoir [START_REF] Johnson | The Ancient Earth[END_REF][START_REF] Li | Biologically recycled continental iron is a major component in banded iron formations[END_REF] (see Supplementary Materials for a detailed discussion). In either case, these Fe isotope data are compatible with a scenario in which near-complete Fe(II) oxidation is occurring at the redox boundary of a stratified water column [START_REF] Busigny | Iron isotopes in an Archean ocean analogue[END_REF] . This suggests that by Stage 4, substantial O2 had started to accumulate in the surface ocean, permitting the development of a stable redoxcline separating deep anoxic and shallow oxygenated seawater. This is consistent with an increasingly important role of free O2 via oxygenic photosynthesis in oxidizing Fe(II). In addition, there is also a decline in the magnitude of positive europium (Eu) anomalies, a high-temperature hydrothermal indicator of Fe input, during Stage 4 IFs (Fig. S3). When combined with increasing oxidation, wanning hydrothermal Fe input could have contributed to falling Fe(II) levels in seawater, which in turn, would have alleviated Fe toxicity issues for cyanobacteria [START_REF] Swanner | Modulation of oxygen production in Archaean oceans by episodes of Fe (II) toxicity[END_REF] . Collectively, these events would accelerate the expansion and proliferation of cyanobacteria [START_REF] Jones | Iron oxides, divalent cations, silica, and the early earth phosphorus crisis[END_REF] , finally giving rise to well-oxygenated surface seawater. Greater O2 availability could also have benefited microaerophilic and nitrate dependent Fe(II)-oxidizers [START_REF] Melton | The interplay of microbially mediated and abiotic reactions in the biogeochemical Fe cycle[END_REF] . Further, a progressive and pervasively oxygenated surface seawater layer would also imply some degree of free O2 accumulation in the atmosphere assuming gas-water exchange equilibrium. Such equilibrium would be fully consistent with geochemical proxy records for oxidative weathering (e.g., chromium abundance, S isotopes) in rocks since ~2.5 Ga [START_REF] Warke | The Great Oxidation Event preceded a Paleoproterozoic "snowball Earth[END_REF][START_REF] Konhauser | Aerobic bacterial pyrite oxidation and acid rock drainage during the Great Oxidation Event[END_REF][START_REF] Philippot | Globally asynchronous sulphur isotope signals require re-definition of the Great Oxidation Event[END_REF] . At this stage MIF-S signals are still preserved in sedimentary sulfides, although there is an intriguing possibility that the MIF-S signal was intermittent during this period [START_REF] Guo | Reconstructing Earth's surface oxidation across the Archean-Proterozoic transition[END_REF][START_REF] Warke | The Great Oxidation Event preceded a Paleoproterozoic "snowball Earth[END_REF][START_REF] Philippot | Globally asynchronous sulphur isotope signals require re-definition of the Great Oxidation Event[END_REF] .

Stage 5 corresponds with the permanent loss of MIF-S signals between 2.4 and 2.3 Ga, which effectively constrains the full expression of the GOE [START_REF] Guo | Reconstructing Earth's surface oxidation across the Archean-Proterozoic transition[END_REF][START_REF] Luo | Rapid oxygenation of Earth's atmosphere 2.33 billion years ago[END_REF][START_REF] Bekker | Dating the rise of atmospheric oxygen[END_REF] . This stage is defined by rare IF deposition and low Mn/Fe ratios (Fig. 1b). We suggest that this reflects the transition from Fe-rich to Fe-poor oceans. This could be linked to either the large-scale oxidative drawdown of the marine Fe(II) reservoir as evidenced by the burial of the giant IF deposits of South Africa and Western Australia, and Fe isotope systematics indicative of Rayleigh fractionation of a significantly diminished Fe(II) pool at the end of Stage 4, or waning hydrothermal Fe(II) input, as shown by weakly positive Eu anomalies observed in the Stage 5 IFs (Fig. S3). The studied IFs and ironstones, for the first time, display a narrow and restricted range of δ [START_REF] Rouxel | Iron isotope constraints on the Archean and Paleoproterozoic ocean redox state[END_REF] Fe values, which are indistinguishable from those of bulk silicate Earth (-0.5 < δ 56 Fe < 0.3 ‰) [START_REF] Johnson | The Ancient Earth[END_REF][START_REF] Beard | Iron isotope constraints on Fe cycling and mass balance in oxygenated Earth oceans[END_REF][START_REF] Severmann | The effect of plume processes on the Fe isotope composition of hydrothermally derived Fe in the deep ocean as inferred from the Rainbow vent site, Mid-Atlantic Ridge, 36 14′[END_REF] . This observation is an expected consequence of essentially quantitative Fe(II) oxidation within a well-oxygenated surface ocean in gas exchange equilibrium with the high-O2 atmosphere, consistent with other independent estimates for atmospheric pO2 [START_REF] Bekker | Oxygen overshoot and recovery during the early Paleoproterozoic[END_REF] .

Modeling ocean-atmosphere O2 abundance

In order to quantitatively estimate the dissolved O2 concentration of the surface ocean and corresponding atmospheric pO2, we focused on shallow-marine GIFs and ironstones, which feature Fe ooids. In contrast to deep-marine BIFs, these facies were deposited in water depths close to, or above, storm and fair-weather wave base in near-shore environments [START_REF] Bekker | Iron Formation: The Sedimentary Product of a Complex Interplay among Mantle, Tectonic, Oceanic, and Biospheric Processes[END_REF][START_REF] Konhauser | Iron formations: A global record of Neoarchaean to Palaeoproterozoic environmental history[END_REF] . This is important because unlike deep-marine settings, shallow seawater is unlikely to be deficient in dissolved O2 relative to the overlying oxygenated atmosphere. These O2 oases likely included the 3.2 Ga Fig Tree and Moodies groups 39, 40 , the 2.95 Ga Sinqeni Formation 11 , and the 2.94 Ga Red Lake Greenstone Belt, and would have been characterized by shallow oxygenated waters set against a backdrop of prevailing anoxic conditions in both the water column and atmosphere.

While oases may have been transient in nature, leading to the generation of 'whiffs' of O2 9 , modelling has also suggested that they may have been more static and laterally extensive, correlated to areas of high nutrient input or upwelling coastal waters [START_REF] Olson | Quantifying the areal extent and dissolved oxygen concentrations of Archean oxygen oases[END_REF] . In either event, these oases would have been characterized by low dissolved O2 concentrations on the order of 1-10 µmoll/kg 53 , which could have supported the oxidation of Fe(II) and correspondent deposition of IFs.

The first appearance of GIFs in the geological record is marked by the ~2.95 Ga Nconga Formation, and followed by the ~2.45 Ga Griquatown and ~2.31 Ga Timeball Hill formations, the latter two spanning a critical interval pertinent to the GOE. The Fe isotope composition of these facies show a progression through time from positive δ 56 Fe values, to a mixture of both positive and negative values, to values negligibly fractionated from bulk silicate Earth (Fig. 1a). Therefore, these facies record a marked shift from partial to quantitative Fe(II) oxidation in shallow seawater around ~2.4 Ga. Assuming gas exchange equilibrium and given that nearly all of Fe in the latter two facies were oxidized ultimately by free O2, partial Fe(II) oxidation in shallow seawater represents robust evidence for low atmospheric pO2, whereas quantitative Fe(II) oxidation indicates comparatively higher pO2. We can use a Rayleigh distillation model (see Supplementary Materials) to invert these Fe isotope data for an estimate of the extent of Fe(II) oxidation (Fig. 2a). The results show a statistically significant difference between the Griquatown GIF and Timeball Hill ironstone, of which the latter exhibits δ [START_REF] Rouxel | Iron isotope constraints on the Archean and Paleoproterozoic ocean redox state[END_REF] Fe values that we consider as clear evidence for complete Fe(II) oxidation (Fig. 2a). We can then combine these observations with a statistical model of Fe(II) oxidation kinetics taking into account a wide range of possible marine pH values, surface water temperatures, and variable residence times of water masses within the shallow ocean (see Supplementary Materials). We find that the Fe isotope signatures observed in the Timeball Hill ironstone are extremely difficult to explain unless the dissolved O2 concentration of their contemporaneous shallow seawater was above 5 µmol/kg, corresponding to atmospheric pO2 above ~1-2% PAL. Further, partial Fe(II) oxidation responsible for the deposition of the Griquatown GIF most likely requires dissolved O2 and atmospheric pO2 below 1 µmol/kg and ~0.1-0.5% PAL, respectively (Fig. 2b). We also invert the Fe isotope data of those BIFs deposited in Archean-early Paleoproterozoic ocean oases for estimating dissolved O2 concentration of the photic zone (Fig. 2 andS14). We show that dissolved O2 was below 1 µmol/kg except for the ~2. 34 Ga photic zone where levels rose above 5 µmol/kg O2.

Our results strongly support a dominantly anoxic world before ~2.4 Ga with local oxygen oases containing < 1 µmol/kg O2, broadly consistent with previous modeling results for a typical O2 concentration (1-10 µmol/kg O2) within an Archean oasis [START_REF] Olson | Quantifying the areal extent and dissolved oxygen concentrations of Archean oxygen oases[END_REF][START_REF] Reinhard | Oxidative sulfide dissolution on the early Earth[END_REF] . Given that a dissolved O2 concentration of 1 µmol/kg is equivalent to a roughly 500-fold oversaturation with respect to an atmospheric pO2 of ~10 -5 PAL (Fig. 2b), it is reasonable to imagine severe ocean-atmosphere disequilibrium with respect to O2 on a reducing Archean Earth surface within biologically productive regions of the ocean. Assuming that the carbon export flux from the photic zone was dominated by oxygenic photosynthesis in local oases, such a low O2 state necessitates that net primary productivity be significantly lower than that of equivalent modern environments, especially given the expectation of increased carbon burial efficiency during the Archean [START_REF] Laakso | Limitations on limitation[END_REF] .

As the export flux can also be expected to respond to rates of primary productivity, which are in turn limited by the recharge of nutrients (e.g., phosphorus (P) and N), low dissolved O2 concentrations require sustained low nutrient conditions prior to ~2.4 Ga. For instance, mass balance considerations have shown that for the Proterozoic, depressed net primary productivity, driven by P limitation, is necessary to maintain a low oxygen state [START_REF] Laakso | Limitations on limitation[END_REF] ; there is no reason to expect that this should be different prior to the GOE. If P is considered to be the nutrient ultimately controlling marine primary productivity in the Archean and early Paleoproterozoic oceans [START_REF] Tyrrell | The relative influences of nitrogen and phosphorus on oceanic primary production[END_REF] , there should be a clear difference with respect to the oceanic inventory of P before and after ~2.4 Ga. A subtle distinction in this regard, is the difference between the overall size of the P reservoir through time and how much is available or accessible to the biosphere. While IFs have provided evidence for a sizeable P reservoir 67 , the record of shallow marine environments stands in stark contrast suggesting that P limitation characterized much of Earth's history [START_REF] Reinhard | Evolution of the global phosphorus cycle[END_REF] . This difference may reflect the enhanced sequestration of P by Fe-phases (oxides, green rusts, viviantie) [START_REF] Reinhard | Evolution of the global phosphorus cycle[END_REF] , or the limited recycling of P during biomass remineralization [START_REF] Kipp | Biomass recycling and Earth's early phosphorus cycle[END_REF] . In either event, this would have limited the P available to primary producers in shallow waters and imposed a direct effect on the degree of primary productivity and, by extension, O2 production. Based on previous Earth system model simulations [START_REF] Olson | Quantifying the areal extent and dissolved oxygen concentrations of Archean oxygen oases[END_REF] , the predominantly anoxic surface ocean implies that the oceanic P inventory of the Archean may have been as low as 25% of the modern, whereas P levels similar to, or higher than, the modern may have resulted in the establishment of a well-oxygenated surface ocean at ~2.31 Ga.

Conclusions

In sum, our findings provide new and complementary insights into how the redox structure of the ocean and atmosphere varied through the Archean to early Paleoproterozoic (Fig. 3). Coupled Fe isotope and Mn/Fe ratio data in IFs extend the first production and local accumulation of O2 in the ocean to at least the latest Paleoarchean, and, when viewed in light of MIF-S signals [START_REF] Busigny | Iron and sulfur isotope constraints on redox conditions associated with the 3.2 Ga barite deposits of the Mapepe Formation (Barberton Greenstone Belt, South Africa)[END_REF] , further confirm that oxygenic photosynthesis was active well before the first permanent oxygenation of the atmosphere during the GOE and that free O2 produced via oxygenic photosynthesis started to play a role in regulating the oxidation of marine Fe(II) since ~3.3 Ga. However, despite the early presence of cyanobacteria, they probably had no critical direct effect on marine Fe cycling initially, and photoferrotrophy remained the dominant mechanism for Fe(II) oxidation in the Archean ocean, as oxygenated shallow seawater oases in the Archean were transient and spatially restricted.

Importantly, our results argue for a pervasive oxygenation of the shallow ocean during the GOE, with dissolved O2 concentrations in surface ocean increasing from below 1 to at least 5 µmol/kg. This clearly marks the revolutionary and progressive transition of the dominant oxidative mechanism of marine Fe(II) from photoferrotrophy to free O2 produced via cyanobacteria at this time. Such a transition could have been closely linked to the growing degree of net primary productivity in the marine environments, ultimately contributing to the progressive, and permanent, oxidation of Earth's surface oceans and atmosphere.

Materials and Methods

IF samples were collected from surface outcrop and drill holes and prepared as polished thin sections and analyzed using reflected light microscopy, Raman spectroscopy, and scanning electron microscopy. Wellpreserved IF samples were crushed in an agate mill to a fine powder, and sample powders were digested for bulk rock geochemical analyses. Major and trace element compositions were determined using an

FiguresFig. 1 .

 1 Figures

Fig. 2 .

 2 Fig. 2. A model for Fe(II) oxidation in the ~3.24 Ga Moodies BIF, ~2.95 Ga Sinqeni BIF, ~2.46 Ga BIFs in western Australia and South Africa, ~2.45 Ga Griquatown GIF, and ~2.31 Ga Timeball Hill ironstone as a function of dissolved O2 abundance of the surface ocean or photic zone. Shown in (a) are estimates of the fraction of Fe(II) oxidized in shallow marine environments (fox) based on a Rayleigh distillation model assuming two overall isotope effects (ɑ = 1.0 (left) and 2.0‰ (right)) for the combined processes of Fe(II) oxidation and precipitation as Fe(III)-(oxyhydr)oxides (see Supplementary Materials). Results are shown for the combined data (filled circles) and data filtered to exclude non-fractionated samples (open circles). Error bars show ±1σ for the entire subsampled dataset. Shown in (b) are results from a kinetic model of Fe(II) oxidation, resampled 10,000 times at each atmospheric pO2 value (relative to PAL) across a range of seawater pH and temperature values. Note that every pO2 value corresponds to one shallow seawater dissolved O2 concentration (µmol/kg) assuming gas-water exchange equilibrium. Error bars show ±1σ. Also note that partial Fe oxidation during the Archean is indicative of low net primary productivity. The systematics of the observed ~2.45 Ga samples are difficult to explain unless dissolved O2 and atmospheric pO2 were below 1 µmol/kg and ~0.1 % PAL, respectively, while the ~2.31 Ga Timeball Hill data imply dissolved O2 at least 5 µmol/kg and a minimum atmospheric pO2 of ~1% PAL during GOE.

Fig. 3 .

 3 Fig. 3. Schematic diagram showing the temporal progression of oxygenation of the Archean to early Paleoproterozoic ocean and atmosphere.

  . Accordingly, the oxidation of Fe(II) that facilitated IF genesis during Stage 1 was likely photoferrotrophy 24 . Compared to Stage 1, IFs deposited during Stage 2 (3.3 to 2.8 Ga) are characterized by lower δ

inductively coupled plasma-mass spectrometer (ICP-MS). For Fe isotope analysis, Fe was purified from sample solutions using ion-exchange chromatography and measured using a multi-collector ICP-MS. Refer to Supplementary Materials for detailed methods. All data discussed in the paper are included in Supplementary Tables S1-S2.

ACKNOWLEDGMENTS:

The manuscript benefitted from discussions with Andrey Bekker. We also acknowledge Xuefang Wu for helping to measure Fe isotopes and Andrey Bekker for providing the Timeball


M Levier

M Roy-Barman

L Foliot

A Dapoigny

F Lacan

Distribution of Pa in the Atlantic Sector of the Southern Ocean: Tracking scavenging during water mass mixing along neutral density surfaces

Keywords: Protactinium, Mass spectrometry, marine particles, Southern Ocean, isopycnal mixing

    Distribution of Pa in the Atlantic sector of the Southern Ocean: Tracking scavenging during water mass mixing along neutral density surfaces

Introduction

Protactinium 231 ( 231 Pa) is a radioactive isotope produced at a constant rate in seawater by the decay of uranium 235 ( 235 U). Dissolved 231 Pa has a residence time in seawater of a few centuries controlled by scavenging by marine particles, so most of the 231 Pa produced in situ is transported toward the sediment. 230 Th is also a particle reactive isotope produced at constant rate in seawater by the decay of uranium 234 ( 234 U). The difference of particle reactivity of 231 Pa and 230 Th induces deviations of the 231 Pa/ 230 Th ratios of seawater, marine particles and sediments from the production activity ratio (0.093) (Rutgers van der [START_REF] Van Der Loeff | On the warm inflow at the eastern boundary of the Weddell Gyre[END_REF]. The Pa/Th ratio in sediments is used as a paleoproxy of the ventilation of deep water in the meridional circulation of the Atlantic Ocean.

Along a meridional transect of the Atlantic Ocean, 231 Pa/ 230 Th ratio measured in seawater increases from north to south. South of the Atlantic Ocean, a change of the 231 Pa/ 230 Th ratio is observed due to the enhanced scavenging of 231 Pa by diatoms [START_REF] Scholten | Advection and scavenging: Effects on 230Th and 231Pa distribution off Southwest Africa[END_REF]. While the affinity of 231 Pa for biogenic opal is well recognized [START_REF] Chase | The influence of particle composition and particle flux on scavenging of Th, Pa and Be in the ocean[END_REF][START_REF] Demaster | The supply and accumulation of silica in the marine environment[END_REF][START_REF] Gdaniec | 231 Pa and 230 Th in the Arctic Ocean: Implications for boundary scavenging and 231Pa230Th fractionation in the Eurasian Basin[END_REF][START_REF] Geibert | Adsorption of thorium and protactinium onto different particle types: experimental findings 11Associate editor: S. Krishnaswami[END_REF][START_REF] Lin | Adsorption and fractionation of thorium and protactinium on nanoparticles in seawater[END_REF][START_REF] Yu | Similar rates of modern and last-glacial ocean thermohaline circulation inferred from radiochemical data[END_REF], there is no direct quantification of the affinity of 231 Pa for biogenic opal in suspended particles from the water column [START_REF] Hayes | Intensity of Th and Pa scavenging partitioned by particle chemistry in the North Atlantic Ocean[END_REF] and on the impact of biogenic opal dissolution in deep waters on the 231 Pa particulate flux [START_REF] Gdaniec | 231 Pa and 230 Th in the Arctic Ocean: Implications for boundary scavenging and 231Pa230Th fractionation in the Eurasian Basin[END_REF]. 231 Pa budget of the Atlantic sector of the Southern Ocean shows that the main source of 231 Pa to the Antarctic Circumpolar Current (ACC) is the inflow from the Atlantic Ocean, half of it being removed by scavenging in the Weddell gyre (Rutgers van der Loeff et al., 2016). The other half of the inflowing Pa is exported to the Pacific sector of the Southern Ocean where 231 Pa is also scavenged or exported to the Pacific Ocean through isopycnal mixing [START_REF] Pavia | Isopycnal Transport and Scavenging of 230 Th and 231 Pa in the Pacific Southern Ocean[END_REF]. The removal of 231 Atlantic Ocean. The Bonus GoodHope section samples were previously analyzed for thorium isotopes [START_REF] Roy-Barman | Thorium isotopes in the Southeast Atlantic Ocean: Tracking scavenging during water mass mixing along neutral density surfaces[END_REF].

Thorium isotopes transport across the ACC was described with an advection-diffusionscavenging model along isopycnal surfaces, allowing to estimate isopycnal advection and eddy diffusion coefficients as well as particles settling velocity. We want to extend this model to protactinium isotopes to get further insight in the behavior of 231 Pa, notably to determine mixing rate of Pa across the ACC. By comparing 231 Pa and Si data, we want to constrain the processes controlling 231 Pa scavenging. Finally, 230 Th and 231 Pa data comparison will give insight in the fractionation of these elements by marine particles. The main water masses and currents encounter during the Bonus GoodHope cruise were Final published version here : https://doi.org/10.1016/j.dsr.2022.103951 described in details elsewhere [START_REF] Abadie | Iron isotopes reveal distinct dissolved iron sources and pathways in the intermediate versus deep Southern Ocean[END_REF][START_REF] Bown | The biogeochemical cycle of dissolved cobalt in the Atlantic and the Southern Ocean south off the coast of South Africa[END_REF][START_REF] Chever | Physical speciation of iron in the Atlantic sector of the Southern Ocean along a transect from the subtropical domain to the Weddell Sea Gyre[END_REF][START_REF] Garcia-Solsona | Rare earth elements and Nd isotopes tracing water mass mixing and particle-seawater interactions in the SE Atlantic[END_REF]. The section is characterized by three main hydrological and circulation domains: the Subtropical domain (Station S1). the Antarctic Circumpolar Current (ACC; S2 to S4) and the Weddell Sea Gyre (S5). The subtropical domain is bounded to the south by the subtropical front (STF), which was located around 41°S-42°S, north of S2. The ACC domain includes the Subantarctic Front (SAF; ~45°S) and the Polar Front (PF; ~50°S) (Fig. 1). It is bounded to the south by the southern ACC boundary (Sby). From the Sby to the Antarctica, water circulation is dominated by the cyclonic Weddell Gyre (WG; S5) [START_REF] Roy-Barman | Thorium isotopes in the Southeast Atlantic Ocean: Tracking scavenging during water mass mixing along neutral density surfaces[END_REF].

Methods

Hydrologic context

Water mass transport along the BGH section is mainly zonal. In the subtropical domain (station S1), it is dominated by the westward flowing Agulhas Current in the upper water column (z < 1500 m) and at depth by the eastward transport of SE-NADW [START_REF] Sokolov | Circumpolar structure and distribution of the Antarctic Circumpolar Current fronts: 1. Mean circumpolar paths[END_REF].

Stations S2, S3 and S4 are located within the Antarctic Circumpolar Current, flowing eastward throughout the water column. Finally, the Weddell gyre is a mainly wind-driven cyclonic gyre [START_REF] Van Der Loeff | On the warm inflow at the eastern boundary of the Weddell Gyre[END_REF], constrained by topographic boundaries except in the East, where mixing occurs between ACC and Weddell Gyre water masses [START_REF] Gouretski | Late slowdown of the Atlantic Meridional Overturning Circulation during the Last Glacial Inception: New constraints from sedimentary (231Pa/230Th)[END_REF][START_REF] Schröder | On the structure and the transport of the eastern Weddell Gyre[END_REF]. The Circumpolar Deep Water (CDW) inflowing the Weddell Gyre, is called Warm Deep Water (WDW). It circulates from the East to the West in the southern part of the Weddell Gyre. The WDW is modified by upwelling into and mixing with shallower water and makes its way to the southern part of the Weddell Sea. By cooling and brine rejection during winter, the cold Weddell Sea Deep Water (WSDW) and Weddell Sea Bottom Water (WSBW) are formed [START_REF] Nicholls | Ice-ocean processes over the continental shelf of the southern Weddell Sea, Antarctica: A review[END_REF] and flow northward, following the cyclonic gyre circulation.

A significant fraction of the WSDW is exported to the North and mixes with the ACC at the Scotia Ridge. The remaining fraction flows toward the northeast of the Weddell Gyre, where the station S5 is located. This water partly recirculates in the gyre by mixing with the Circumpolar Deep Water incoming from east to form the new WDW [START_REF] Deacon | The Weddell gyre[END_REF][START_REF] Gouretski | Late slowdown of the Atlantic Meridional Overturning Circulation during the Last Glacial Inception: New constraints from sedimentary (231Pa/230Th)[END_REF].

The BGH section is approximately perpendicular to the mean flows of the ACC, so the BGH stations are not hydrologically connected and there might be no direct mixing between the stations. However, slower meridional transport were noted on meridional sections of water mass tracers such as salinity, temperature or iron isotopes, revealing notably the northward transport of AAIW and AABW, and southward transport of NADW [START_REF] Abadie | Iron isotopes reveal distinct dissolved iron sources and pathways in the intermediate versus deep Southern Ocean[END_REF]Roy-Barman Submitted, non-definitive, version. Final published version here : https://doi.org/10.1016/j.dsr.2022.103951 et al., 2019). This meridional transport is assumed to be a residual flow from the Atlantic Ocean circulation in which lateral eddy fluxes largely balance the wind-driven circulation [START_REF] Marshall | Closure of the meridional overturning circulation through Southern Ocean upwelling[END_REF]. In the following, we apply a simple transport model of advection-diffusionscavenging to link tracer properties measured along the BGH meridional section [START_REF] Roy-Barman | Thorium isotopes in the Southeast Atlantic Ocean: Tracking scavenging during water mass mixing along neutral density surfaces[END_REF]. The transports are not assumed to be a true meridional mixing along the BGH section, but it allows to estimate the residual meridional component of the mixing (advection, diffusion) embedded in the dominant zonal of the ACC Atlantic sector.

Sampling

The BONUS GoodHope cruise occurred from February 8th 2008 to March 17th 2008, during the late austral summer, on board Marion-Dufresne II Research Vessel. Samples were collected along a transect between 34°S 19°E and 51°S 0°W and then along the Greenwich meridian from 51°S to 58°S. Five full depth vertical profiles were sampled (S1-S5 stations) for both dissolved and particulate 231 Pa analysis. Three depths were collected for dissolved 231 Pa analysis at an intercalibration station. Seawater samples were collected with Niskin bottles mounted on a rosette, equipped with a CTD sensor. Challenger Oceanic in situ pumps (ISP) were used to filter large seawater volumes through SUPOR filters (pore size: 0.45 µm, filter diameter: 293 mm) seawater volumes filtered were ranging from 200 L to 1000 L.

Analytical procedure

Filtered seawater

Ten liters of filtered seawater (Nuclepore TM , 90mm diameter, 0.4 μm pore size) were acidified at pH = 2 with HCl on board. The samples were processed in a clean lab for the simultaneous extraction of 227 Ac and 231 Pa as describe in [START_REF] Levier | Determination of low level of actinium 227 in seawater and freshwater by isotope dilution and mass spectrometry[END_REF]. To summarize, samples were spiked with 233 Pa, previously milked from 237 Np on hydrated silica (Guihou et al., 2010). Pa isotopes were preconcentrated using manganese oxides co-precipitation [START_REF] Ghaleb | Improved method for radium extraction from environmental samples and its analysis by thermal ionization mass spectrometry[END_REF][START_REF] Van Der Loeff | Determination of natural radioactive tracers[END_REF]. The precipitate was recovered on NucleoporeTM filter (diameter 142mm, 0.45 µm pore size). The precipitate still on the filter was dissolved in a bath of 6 M HCl with 100 µL of H2O2 and 100 µL of 27 M HF during 30min.

The filter was then rinsed with 6M HCl. After evaporation of the dissolution bath, the residue was dissolved in 0.5 mL of 9 M HCl and loaded on an anion-exchange column (AG1-X8 resin, 200-400 mesh) to separate the protactinium fraction, eluted with 9M HCl + 0.26M HF, from the major elements, uranium, thorium and rare earth elements (REEs) and actinium [START_REF] Levier | Determination of low level of actinium 227 in seawater and freshwater by isotope dilution and mass spectrometry[END_REF].

Particulate samples

Protactinium from particulate samples was recovered and purified during same chemistry as the particulate thorium isotopes [START_REF] Roy-Barman | Thorium isotopes in the Southeast Atlantic Ocean: Tracking scavenging during water mass mixing along neutral density surfaces[END_REF]. Filters from the in situ pumps were cut into pieces using ceramic scissors. Several leaching steps were necessary to recover Th without attacking the filters. Filter pieces were first leached with 200 mL of 6N HCl and 1.5 mL of 24-25N HF in a 300 mL Teflon beaker for 2 days at 75°C. The filter pieces were removed from the leaching solution and rinsed with a MQ water squeeze bottle over the leaching beaker.

The filter pieces were saved apart. The leaching and rinsing solutions were evaporated down to ~10 mL and transferred into a 30 mL Teflon beaker. Then, the filter pieces were leached a second time with 150 mL of 7N HNO3 and 0.05 mL of 29N HF for 2 days at 75°C. The filter pieces were removed from the leaching solution and rinsed with MQ water that was again recovered into the leaching beaker. After evaporation to a few mL, this solution was added to the first leaching and rinsing solution. The 300 mL Teflon beaker was rinsed in warm diluted HNO3 to remove any particle sticking on the beaker walls and the resulting solution was also added to the 30 mL beaker. The resulting solution was then spiked with 233 Pa (and 229 Th ), dried and dissolved again in a solution with 4 mL of 14N HNO3and 1 mL of 12N HCl. After 1 night on a hot plate, the solution was dried and the residue was dissolved again in 10 ml of 1N HNO3.

Since the filters were not rinsed immediately after filtration with distilled water on board, it was preferable to remove the salt before the column chemistry. Therefore, 40µL of a Fe solution (60 mg/g) were added. After 1h-heating, Fe was precipitated by raising the pH to about 8 with NH3. The Fe precipitate (that coprecipitates Th and Pa isotopes) was separated by centrifugation and rinsed several times. Finally, it was dissolved in 0.25 ml of 8N HNO3, ready for loading on an anionic column. Pa and Th isotopes were separated from Fe by ion exchange chromatography on a small volume (0.5 ml) column of AG1X8 (200-400 mesh) resin [START_REF] Gdaniec | Thorium and protactinium isotopes as tracers of marine particle fluxes and deep water circulation in the Mediterranean Sea[END_REF], adapted from Jeandel et al., 2011). Every analysis was performed on a Multiple Collection Inductively Coupled Plasma Mass Spectrometer (MC-ICPMS) Neptune plus (Thermo-Fischer) mounted with a jet interface and an Aridus II desolvatation system based on a protocol derived from Guihou et al., 2010. The acquisition conditions of Pa measurements are given in [START_REF] Gdaniec | Thorium and protactinium isotopes as tracers of marine particle fluxes and deep water circulation in the Mediterranean Sea[END_REF]. The calibration of 233 Pa spike solution was made against an in-house 231 Pa standard, used for intercalibration exercises ( Gdaniec et al., 2019 ;Gdaniec et al., in prep). Particle samples from stations S1 and S2, spike calibration was lost during the chemical separation. This loss made impossible the precise correction for the chemical yield. To circumvent this problem, we noted that for station S3, S4 and S5, the chemical yield ranged from 20% to 100%, but for each series the highest yield ranged from 80% to 100%. Therefore, we attributed a chemical yield of 90% ± 10% to the samples of stations S1 and S2 with the highest count rate on 233 Pa and deduced the concentration of the spike solution by using the sensitivity (count per second per ppt) measured on a 231 Pa-233 U calibration solution. This procedure was checked on stations S3, S4 and S5, the concentrations deduced for all samples fall within ± 10% of concentrations deduced from the true spike concentration obtained by the spike calibration against the home standard.

Mass spectrometry

231 Pa concentrations were corrected for ingrowth from 235 U decay between the on-board sampling and the U-Pa chemical separation [START_REF] Levier | Determination of low level of actinium 227 in seawater and freshwater by isotope dilution and mass spectrometry[END_REF]. All uncertainties were propagated through signal processing equation and correction equation, and each of them were expressed with a confidence interval of 95% (2 times standard error: 2σn).

Results

In order to consider exclusively 231 ( 235 U/ 238 U)nat is abundance ratio between natural uranium isotopes, here 1/137.818 mol/mol [START_REF] Livermore | Evaluating the robustness of a consensus 238U/235U value for U-Pb geochronology[END_REF][START_REF] Condon | Isotopic composition ( 238 U/ 235 U) of some commonly used uranium reference materials[END_REF]. The mean crustal activity ratio ( 238 U/ 232 Th)litho = 0.4 ± 0.1, estimated south of the Antarctic Polar Front (Rutgers van der [START_REF] Van Der Loeff | On the warm inflow at the eastern boundary of the Weddell Gyre[END_REF][START_REF] Venchiarutti | Scavenging of 231 Pa and thorium isotopes based on dissolved and size-fractionated particulate distributions at Drake Passage (ANTXXIV-3)[END_REF].

Dissolved protactinium 231 corrected from the lithogenic contribution ( 231 Pad-xs) measured for the Bonus GoodHope cruise ranged from 0.19 ± 0.02 fg/kg in surface water to 3.45 ± 0.06 fg/kg (Table ) in bottom water with a general increase of the concentration with depth (Figure 2). At the stations S1 to S3, we observe a similar profile with a linear increase from the surface to around 1500m. At Station S1, we see a maximum concentration of 3.34 ± 0.06 fg/kg at 3000m, in NADW, and decrease to 2.79 ± 0.06 fg/kg at 4000m in the AABW. At station S3, there is a sharp decrease of the concentration near depth, from 2.75 ± 0.05 fg/kg at 3000m to 1.67 ± 0.04 fg/kg in the bottom water. This decrease was already noted for dissolved REE [START_REF] Garcia-Solsona | Rare earth elements and Nd isotopes tracing water mass mixing and particle-seawater interactions in the SE Atlantic[END_REF] and thorium isotopes [START_REF] Roy-Barman | Thorium isotopes in the Southeast Atlantic Ocean: Tracking scavenging during water mass mixing along neutral density surfaces[END_REF], (possibly due to hydrothermal scavenging, or a rich deep nepheloid layer.)

The Bonus GoodHope cruise occurred 3 weeks after the ZeroDrake cruise, where the Greenwich meridian was also sampled across the ACC for 231 Pa/ 230 Th studies (Rutgers van der Loeff et al., 2016;[START_REF] Venchiarutti | Scavenging of 231 Pa and thorium isotopes based on dissolved and size-fractionated particulate distributions at Drake Passage (ANTXXIV-3)[END_REF]. Thorium isotopes were consistent between the two cruises [START_REF] Roy-Barman | Thorium isotopes in the Southeast Atlantic Ocean: Tracking scavenging during water mass mixing along neutral density surfaces[END_REF]. Consistent data are also obtained for 231 Paxs, when we compare stations S2 to S4 with nearby ZeroDrake stations (Figure ES10). At these stations, the BGH concentrations tend to be on the low side of the Drake values. However, an intercalibration exercise between LSCE and AWI, on strictly identical samples, proved an agreement within a few percent for 231 Pa measured in deep arctic sample (Gdaniec et al in prep).

By contrast, in the Weddell Gyre, the dissolved 231 Paxs was almost twice less concentrated at Station S5 compared to station 131 of the ZeroDrake cruise. There is about 150 km between these stations (0.00E, 57.5S for the BGH sampling point and 0.00E, 59.0S for the ZeroDrake one). Possible reasons for this discrepancy will be discussed in section 4.1. strong difference in 231 Paxs concentration despite being in the same hydrological settings (Figure 4). Difference in 231 Paxs content between these stations could have different origins:

Deep water in the Weddell gyre, WSDW is newly formed deep water in Weddell Sea were reflected by the highly oxygenated signature (about 250 µmol/L). The dissolved oxygen in bottom water of station S5 (O2 = 248.3 µmol/kg, [START_REF] Branellec | Projet GoodHope. Campagne BONUS / GOODHOPE[END_REF] is higher than at ZeroDrake station 131 (O2 = 241.7 µmol/kg) (Fahrbach & Baar, 2010), suggesting a younger water at station S5 (Figure ). Assuming an oxygen utilization rate (OUR) of about 0.12-0.14 µmol/L/y [START_REF] Broecker | Radiocarbon decay and oxygen utilization in the Deep Atlantic Ocean[END_REF][START_REF] Feely | Oxygen Utilization and Organic Carbon Remineralization in the Upper Water Column of the Pacific Ocean[END_REF], it would take about 50 y to produce the observed O2 deficiency of ~7 µmol/L observed at station 131. We note that the difference of FTh/Pa across the ACC range from 0.57 ± 0.06 in shallow water of station S5 to 37.6 ± 16.2 at 1450 m depth at station S2 (Figure 5; Table ES1). There is a latitudinal gradient from the north of the ACC, with FTh/Pa ranging from about 1 in surface waters to about 20 in deep waters, to the south of the polar front with FTh/Pa close to 1 throughout the water column, consistent with previous observations across the ACC [START_REF] Chase | The influence of particle composition and particle flux on scavenging of Th, Pa and Be in the ocean[END_REF][START_REF] Venchiarutti | Scavenging of 231 Pa and thorium isotopes based on dissolved and size-fractionated particulate distributions at Drake Passage (ANTXXIV-3)[END_REF][START_REF] Walter | Enhanced scavenging of 231 Pa relative to +Th in the South Atlantic south of the Polar Front: Implications for the use of the 231 Pa/ 230 Th ratio as a paleoproductivity proxy[END_REF]. The variation of FTh/Pa is mainly constrained by the marine particle composition. In polar areas, most of these variations are attributed to the Pa scavenging that is driven by the biogenic silica (BSi) cycle, because 231 Pa has a high affinity for biogenic silica [START_REF] Chase | The influence of particle composition and particle flux on scavenging of Th, Pa and Be in the ocean[END_REF][START_REF] Gdaniec | 231 Pa and 230 Th in the Arctic Ocean: Implications for boundary scavenging and 231Pa230Th fractionation in the Eurasian Basin[END_REF][START_REF] Walter | Enhanced scavenging of 231 Pa relative to +Th in the South Atlantic south of the Polar Front: Implications for the use of the 231 Pa/ 230 Th ratio as a paleoproductivity proxy[END_REF]. Therefore, we compare the 231 Pa and Si concentrations in dissolved and particulate fractions (Figure 6a).

South of the Polar Front, the low 231 Pad/DSi ratio in shallow waters is due to the preferential removal of 231 Pa relative to the Si, due to particles with high Pa/BSi ratio and the high DSi concentration in these surface waters [START_REF] Sarmiento | High-latitude controls of thermocline nutrients and low latitude biological productivity[END_REF]. In contrast, north of the Polar Front, water are poor in DSi [START_REF] Sarmiento | High-latitude controls of thermocline nutrients and low latitude biological productivity[END_REF] Since the FTh/Pa seems to be mainly impacted by the affinity of 231 Pa for the biogenic silica, we use this new dataset to constrain the partition coefficient of 231 Pa on opal in the suspended particle pool. The partition coefficient is defined as the concentration adsorbed on particles normalized by to the dissolved concentration [START_REF] Hayes | Intensity of Th and Pa scavenging partitioned by particle chemistry in the North Atlantic Ocean[END_REF]:

𝐾 𝑑-𝑏𝑢𝑙𝑘 (𝑋) = 𝑋 𝑎𝑑𝑠 𝑋 𝑑𝑖𝑠𝑠 × 1 𝑆𝑃𝑀 = 𝑋 𝑎𝑑𝑠 𝑆𝑃𝑀 𝑋 𝑑𝑖𝑠𝑠 (3.3)
Where X represents 231 Pa or 230 Th, in the adsorbed fraction (~ particulate concentration) and dissolved fraction, and SPM the mass of suspended particle matter in seawater, that we estimated as the sum of the opal, carbonate, particulate organic matter (POM) and lithogenic matter. The opal mass in particle matter in estimated from the BSi concentration [START_REF] Fripiat | Silicon pool dynamics and biogenic silica export in the Southern Ocean inferred from Si-isotopes[END_REF] with the mean opal molar mass (SiO2,0.4(H2O)). The carbonate (CaCO3) concentration was estimated from the particulate Ca concentration (Table . ES4). POM mass was approximated from the particular organic carbon (POC) [START_REF] Cavagna | Water column distribution and carbon isotopic signal of cholesterol, brassicasterol and particulate organic carbon in the Atlantic sector of the Southern Ocean[END_REF] with POM = 2 × POC [START_REF] Winogradow | Seasonal changes in particulate organic matter (POM) concentrations and properties measured from deep areas of the Baltic Sea[END_REF]. The lithogenic matter was approximated from the 232 Th particulate concentration [START_REF] Roy-Barman | Thorium isotopes in the Southeast Atlantic Ocean: Tracking scavenging during water mass mixing along neutral density surfaces[END_REF] with 232 Th concentration in lithogenic matter is about 10 ppm. Kd-bulk(Pa) and Kd-bulk(Th) were calculated for sample where BSi concentrations were available.

The Kd-bulk(Pa) range from 0.062 to 2.79 ×10 6 g/g, and Kd-bulk(Th) range from 0.212 to 9.92 ×10 6 g/g. 231 Pa concentration is assumed to be driven by the Si biogeochemical cycle, but there is only a weak correlation between the Kd-bulk(Pa) of Bonus particles and the BSi proportion in marine particles (R² = 0.17), and no correlation with the Kd-bulk(Th) (R² = 0.02) (Figure ). The Kd-opal(Pa) can be deduced from Kd-bulk(Pa) as follows:

𝐾 𝑑-𝑜𝑝𝑎𝑙 = (𝐾 𝑑-𝑏𝑢𝑙𝑘 -𝑓 𝑙𝑖𝑡ℎ𝑜 ×𝐾 𝑑-𝑙𝑖𝑡ℎ𝑜 -𝑓 𝐶𝑎𝐶𝑂3 ×𝐾 𝑑-𝐶𝑎𝐶𝑂3 -𝑓 𝑃𝑂𝑀 ×𝐾 𝑑-𝑃𝑂𝑀 ) 𝑓 𝑜𝑝𝑎𝑙 (3.4)
Where flitho, fCaCO3, fopal and fPOM are the mass fractions of each phase in seawater, Kd-litho, Kd-CaCO3 and Kd-POM are the partition coefficient of each phase, with respective values : 2.3 ×10 6 ; 0.9 ×10 6 and 0.2 ×10 6 , determined on North Atlantic particles [START_REF] Hayes | Intensity of Th and Pa scavenging partitioned by particle chemistry in the North Atlantic Ocean[END_REF]. Taking into account only particulate samples containing more than 50% of BSi (samples from opal belt and south of the polar front) to avoid uncertainties amplification for BSi-poor samples, we obtain Kd-opal(Pa) = 1.42 ±0.55 ×10 6 g/g (1 standard error). The same approach applied to the Kdopal(Th) provide no significant results, probably because Kd-opal(Th) is in the same range or lower than the Kd of the other phases.

Kd-opal(Pa) crude estimation on suspended particles are in the upper range of previous estimation made on sediment trap material (Kd-opal(Pa) = 1.0 to 1.4× 10 6 , [START_REF] Chase | The influence of particle composition and particle flux on scavenging of Th, Pa and Be in the ocean[END_REF][START_REF] Li | Controversy over the relationship between major components of sediment-trap materials and the bulk distribution coefficients of 230 Th, 231 Pa, and 10 Be[END_REF][START_REF] Luo | On the importance of opal, carbonate, and lithogenic clays in scavenging and fractionating 230 Th, 231 Pa and 10 Be in the ocean[END_REF], and over the estimation made on SiO2 from diatoms in laboratory experiment (Kd-opal(Pa) = 1.9 × 10 5 , Lin et al., 2014 ; Kd-opal(Pa) = 5.0 × 10 5 , [START_REF] Geibert | Adsorption of thorium and protactinium onto different particle types: experimental findings 11Associate editor: S. Krishnaswami[END_REF]. In situ material, suspended particles or sediment trap, have a larger partition coefficient than laboratory experiment on diatoms culture. This value confirms that opal represents a significant contribution to the 231 Pa scavenging, essentially with the opal dominance in particle matter in Southern Ocean.

Isopycnal transport and scavenging of 231 Pa across the ACC

Across the ACC, 230 Th concentrations were shown to be relatively conservative along isopycnal surfaces [START_REF] Roy-Barman | Thorium isotopes in the Southeast Atlantic Ocean: Tracking scavenging during water mass mixing along neutral density surfaces[END_REF][START_REF] Van Der Loeff | On the warm inflow at the eastern boundary of the Weddell Gyre[END_REF].Thorium isotopes transport along the BGH section was represented with an advection-diffusion-scavenging model along isopycnal surfaces (Figure 7). This allowed linking tracers (θ-S-230 Th-232 Th) and estimating transport parameters [START_REF] Roy-Barman | Thorium isotopes in the Southeast Atlantic Ocean: Tracking scavenging during water mass mixing along neutral density surfaces[END_REF]. The transports are not assumed to be a true meridional mixing along the BGH section, but it allows to estimate the residual meridional component of the mixing (advection, diffusion) embedded in the dominant zonal of the ACC Atlantic sector. In the present study, we report that 231 Pa concentrations show the same behavior than 230 Th, with relatively small gradients along the a given isopycnal surfaces (Figure 8). Therefore, we now use 231 Pa data as additional constrain to the model, test its robustness and the effect of an additional equation on the estimation of transport parameters.

As for Th isotopes, we focus on two isopycnal surfaces defined by γn = 27.865 kg/m 3 and γn = 28.094 kg/m 3 (fig. 7). The first isopycnal surface (γn = kg/m 3 ) corresponds to the mixing of the UCDW at station S1 at around a depth of 1500 m and the D-UCDW at station S4 with the isopycnal surface upwells to about 200m. The second isopycnal surface (γn = 28.094 kg/m 3 ), in the southern part of the ACC, corresponds to the mixing between the upper part of the NADW at the station S3 at a depth of about 2500m and the WDW at station S5, at a depth of about 150m. These 2 isopycnal surfaces were chosen because along the Bonus Good Hope section they best correspond to a 2 water masses mixing as indicated by θ-S data. Where Ki and ui are respectively the cross-stream eddy diffusion coefficient and the crossstream isopycnal advection velocity coefficient. Pd is the in situ production of the radiogenic element and wp the settling speed of particles. Ct and Cp are respectively the total concentration and the particulate concentration of the radiogenic element.

The x-axis is parallel to the isopycnal surface in the meridional direction and oriented northward. The z-axis is oriented perpendicular to the isopycnal surface and hence is almost vertical. Just like in the reversible scavenging model, keeping wp out of the partial derivative requires to assume that wp does not vary with z. The Eq. (3.5) was integrated along the isopycnal surface of the BGH section between the stations corresponding to the end members of water mass mixing (stations for S1 and S4 for γn=27.865 kg/m 3 , from the I-UCDW to the DP-UCDW and S3 and S5 for γn=28.094 kg/m 3 , from the SW-NADW/AABW to the AASW/WW). This is a simplified view of water transport because while water moves across the ACC, it is also rapidly advected eastward by the ACC (S2-S4) and the northern limb of the Weddell Gyre (S5).

We assume that Ki, ui, wp and (dCp/dz) are all constant along x in order to obtain analytical solutions. We note Ct=Ct-cons + ΔCt, where Ct-cons is the concentration of the tracer if it behaves as a conservative mixing (only mixing by advection and diffusion) and ΔCt is the deviation of this tracer due to the radioactive production and scavenging on settling particles. The solution is then: )) (3.7)

Where the suffixes A and B represents stations and x the latitudinal position.

For a conservative tracer such as the salinity (S), the concentration is given by:

𝑆 = 𝑒 𝑢 𝑖 𝐾 𝑖 (𝑥-𝑥 𝐴 ) -1 𝑒 𝑢 𝑖 𝐾 𝑖 (𝑥 𝐵 -𝑥 𝐴 ) -1 (𝑆 𝐵 -𝑆 𝐴 ) + 𝑆 𝐴 (3.8)
This equation allows to determine ui/Ki ≈ -1.6 10 -6 m -1 for both isopycnal surfaces [START_REF] Roy-Barman | Thorium isotopes in the Southeast Atlantic Ocean: Tracking scavenging during water mass mixing along neutral density surfaces[END_REF]. Then equations 3.6 and 3.7 were applied to 230 Th and 232 Th to derive a first estimate of the transport parameters (Table 1, see [START_REF] Roy-Barman | Thorium isotopes in the Southeast Atlantic Ocean: Tracking scavenging during water mass mixing along neutral density surfaces[END_REF] for details of the calculations).

Finally, we applied the equations 3.6 and 3.7 were to 230 Th and 232 Th and 231 Pa data of the BGH cruise to obtain a second set of transport parameters (Table 1)

We use a Monte-Carlo simulation (n=50000) over the 3 model parameters (ui, Ki, wp) with a large dispersion of the their a priori values. Further constrains on ui and Ki arise from the ui/Ki ratio solution of the conservative equation. A priori values of dCp/dt covers all the particulate gradients measured along the isopycnals. We determine the coefficient of determination (R²)

between each mixing curves of the Monte-Carlo simulations and the measured concentrations on the isopycnal surface. We use the mean of the R² on the three isotopes as a selection filter, to keep only parameters values that make converging mixing curves toward the measured concentrations. We arbitrarily put the selection at R² = 0.97 for the first isopycnal surface (γn = 27.865) and R² = 0.75 for the second one (γn = 28.094). This selection factor brings a selection of the most fitting parameters, which we derived their mean value and standard deviation (Tab.

1)

Including 231 Pa in the model, there is no significant change on the average estimate on parameters Ki and ui, but the uncertainties are reduced by a factor 2 to 4 (Table 1). Among the model parameters, the particle settling speed is the most affected by the introduction of the Pa constrains with higher settling speed on the isopycnal surface 28.094 kg/m 3 , with a value of 990 ± 85 m/y against 674 ±250 m/y with Th isotopes alone (Table 1). This change of settling speed is calculated with the stations S3 to S5, located on the southern part of the section, where the vertical particulate flux is mainly driven by diatoms (that scavenge Pa efficiently). Hence, 231 Pa sensitivity to scavenging by the biogenic silica allows a better representation of these fast sinking particles. Conversely, calculation along the isopycnal surface γn = 27.865 kg/m 3 is based more on the northern stations, where lower diatom production does not drive the flux of settling particles. Therefore, including Pa data along this isopycnal surface does not affect the particle settling speed, but significantly reduces the uncertainties on ui, Ki and wp. The advectiondiffusion-scavenging model is able to reproduce the data measured on sample from the BGH cruise, linearly distribute along an isopycnal surface (Figure 9). This suggests that the model processes are the most predominant on the 231 Pa cycle of the ACC and our different hypothesis like the steady state are consistent to explain the observed distribution. kg/m 3 (note that for these 2 isopycnals the mean depth considered is around 1200-1400 m). This reflect that 231 Pa is mainly constrained by the advection and diffusion in Southern Ocean although the scavenging flux measured in the opal belt is lower than in water near the African margin. Electronic Support : For each isopycnal surface, the box is bounded to the North and the south by the station corresponding to the end-members of the mixing (S1 and S4 for γn = 27.865 kg/m 3 ; S3 and S5

Conclusion

for γn = 28.094 kg/m 3 ). The horizontal length scale L is half of the distance between these endmember stations (Table 3-2

  Pa upwelling from the South Atlantic in the Weddell gyre results from a double trap : first by the gyre circulation preventing intermediate 231 Pa-rich waters to flow northward, enhancing the 231 Pa scavenging in the center of the gyre, and second by Pa scavenging in the opal belt (area of high diatom production, around 52°S), when Weddell Sea Deep Water (WSDW) flows out the gyre (Rutgers van der Loeff et al., 2016). Here, we analyze dissolved and particulate 231 Pa in samples collected across the Antarctic Circumpolar Current (ACC) during the Bonus GoodHope cruise in the Southern sector the Submitted, non-definitive, version. Final published version here : https://doi.org/10.1016/j.dsr.2022.103951

Figure 1 :

 1 Figure 1 : Map of the stations from the Bonus GoodHope cruise, with the different front across the Southern Ocean

  Pa concentration produced in situ by the 235 U decay, the 231 Pa concentrations measured by mass spectrometry were corrected from the contribution brought by lithogenic material, where we used 232 Th as a proxy. The excess of 231 Pa was expressed as follows: Where 231 Paxs is the excess of 231 Pa corrected from the lithogenic contribution, 231 Pam and of the 231 Pa and 230 Th.

  Figure 4 : 231 Paxs-total versus neutral density in the Weddell gyre during the Bonus GoodHope and ZeroDrake cruises. Red dot: station S5 BGH; blue triangle: station 131 of ZeroDrake

Figure 5 :

 5 Figure 5 : FTh/Pa fractionation factor profiles. The dissolved concentration was interpolated linearly between two measured point on neutral density basis at depths where particulate concentration was measured. FTh/Pa uncertainties are propagated and expressed in 2𝜎n

  resulting of low BSi production and low Pa removal in shallow waters. In deep waters, the difference of Pad/DSi ratio north and south to the Polar Front results from the different compositions of sinking particles: south of the Polar Front, BSi-rich particles release more DSi than BSi-poor particles sinking north to the Polar Front. The general trend of increasing Pap/BSi ratio with depth (Figure 6b) reflects the -definitive, version. Final published version here : https://doi.org/10.1016/j.dsr.2022.103951 dissolution of BSi as particles sink through the water column. Meanwhile, 231 Pa remains on particles and additional Pa might be scavenged at depth. These observations are consistent with the assumption that the 231 Pa is driven by the silicium cycle and the explanation of the meridional gradient given previously (Rutgers van der Loeff et al., 2016).

Figure 6 :

 6 Figure 6 : (a) Pad/Si ratio in the water column. (b) Pap/BSi ratio in the particulate fraction

  -definitive, version. Final published version here : https://doi.org/10.1016/j.dsr.2022.103951

Figure 7 :Figure 8 :

 78 Figure 7 : Distribution of the 231 Pat (Pad + Pap) concentration from the Bonus GoodHope cruise against the neutral density. The studied isopycnal surfaces (𝛾n = 27.865 kg/m 3 and 𝛾n = 28.094 kg/m 3 ) are highlighted by black lines.

Figure 9 :

 9 Figure 9 : Advection-diffusion modelling of Salinity and 231 PaT compared to data, at the isopycnal surface γn= 27.865 kg/m 3 (a) and at the isopycnal surface γn= 28.094 kg/m 3 (b). Stacked red curves are the fitting curved with the best R² obtained with the Monte Carlo simulation.

  Submitted, non-definitive, version. Final published version here : https://doi.org/10.1016/j.dsr.2022.103951 The distribution of 231 Pa in the Southern Ocean is mainly constrained isopycnal advection and diffusion. The crude estimation shows that most of the 231 Pa transit through the advection and diffusion fluxes and few of the 231 Pa is scavenged in the ACC, north to the Polar Front. This makes the 231 Pa a good tracer of the mixing in the Southern Ocean and also a chronometer of this mixing, like in Weddell Gyre where we were able to made an estimation of the age difference between the inner and the peripheral Weddell Gyre water. This new dataset of 231 Pa in this high latitude area also allowed to propose a new estimate of the 231 Pa adsorption coefficient on opal particle fraction which is higher than previous estimations. The use of Pa and Th in the model allowed determining the settling rate of the particles, but large uncertainties remained despite the addition of 231 Pa tracer in the advection diffusion scavenging model. The 2 end-member mixing model presented here could be used only for 2 isopycnal surfaces. More complex mixings could be treated through optimum multiparameter (OMP) analysis (Le Roy et al., 2018) when more 231 Pa data will be available. Another improvement of the model would be to use a variable particulate flux along the isopycnal surface γn=28.094 kg/kg, where the particulate Pa gradient increases by two order of magnitudes between station S3 (10 -6 fg/kg/m) and station S5 (10 -4 fg/kg/m) by crossing the Polar Front, reflecting the change of the particulate phase carrying the 231 Pa scavenging. Finally, a last challenge would be to include other mixing tracers in the model such as Nd isotopes for water mass tracing or 227 Ra or 227 Ac as short-lived chronometers.

Figure

  Figure ES10 : 231 Paxs-dissolved versus neutral during the Bonus GoodHope and ZeroDrake cruises at hydrologically close stations.

  ). The mean height of the box is the depth of the isopycnal halfway to the 2 end-member stations. The width of the box used for an intermediate calculation, noted l, does not appear in the final result. We choose the mean dissolved and particulate of 231 Pa concentration of 231 Pa and 231 Pa at the depth of the isopycnal surface halfway between the 2 end-member stations. The in situ 231 Pa production rate in seawater is P = 25 fg/m 3 /y using the output parameters from the advection-diffusion-scavenging model from

  The cyclonic circulation in the Weddell isolates an inner zone with high231 Pad-xs concentrations (up to 5.50 ±0.31 fg/kg, at station 131 of Rutgers van der Loeff, 2016) whereas the rim has lower concentrations (about 4.14 ±0.51 fg/kg, at station 161 of Rutgers van der Loeff, 2016). The reasons for this concentration gradient are unclear. The low dissolved 231 Paxs measured at Station S5, is in line with the peripheral location of the station in the gyre at the boundary between the Weddell gyre and the southern area of the ACC. This region has a strong gradients, with a enriched concentration in Weddell gyre compare to the ACC, for several radioactive isotopes like 230 Th, 232 Th,

			Submitted, non-definitive, version. Final published version here : https://doi.org/10.1016/j.dsr.2022.103951
		0	0	0	0 0.1	1 0.2	2 231 Pa d-xs (fg/kg) 0.3	3 0.4	4
		500						Station S1
		1000						Station S2
				1000				Station S3
		1500						Station S4
	Depth (m)	3500 2000 2500 3000		2000 3000 Depth (m)	Station S2 Station S1	Station S5
		4000			Station
		4500		4000	S3
					Station
		5000			S4
				5000

Submitted, non-definitive, version. Final published version here : https://doi.org/10.1016/j.dsr.2022.103951 Figure 2 : 231 Paxs concentration in the dissolved phase of the super stations of the Bonus GoodHope cruise

The 231 Pa concentration in excess on the particulate fraction ( 231 Pap-xs) ranged from 0.00081 ± 0.00016 fg/kg in surface water (station S4) to 0.341 ± 0.054 fg/kg in bottom water of station S1 (Figure

3

). The 231 Pap-xs concentration are rather constant through the water column with a sharp increase near the seafloor, except at station S3, where the deepest concentration remains similar to the intermediate waters (Table

)

. Submitted, non-definitive, version. Final published version here : https://doi.org/10.1016/j.dsr.2022.103951 Figure 3 : 231 Paxs concentration in the particulate phase of the super stations of the Bonus GoodHope cruise 4. Discussion 4.1. Dissolved 231 Pa in the Weddell Gyre Dissolved 231 Pa profiles are variable in the Weddell Gyre (Rutgers van der Loeff et al., 2016). 231 Pa (Rutgers van der Loeff

et al., 2016;[START_REF] Venchiarutti | Scavenging of 231 Pa and thorium isotopes based on dissolved and size-fractionated particulate distributions at Drake Passage (ANTXXIV-3)[END_REF]

or Pb

[START_REF] Boye | Distributions of dissolved trace metals (Cd, Cu, Mn, Pb, Ag) in the southeastern Atlantic and the Southern Ocean[END_REF]

. Station S5 and station PS63-161 of ZeroDrake cruise located 150 km further south show a 231 Pa p-xs (fg/kg)

  Submitted, non-definitive, version. Final published version here : https://doi.org/10.1016/j.dsr.2022.103951

  Submitted, non-definitive, version. Final published version here : https://doi.org/10.1016/j.dsr.2022.103951

		𝐶 𝑡-𝑐𝑜𝑛𝑐 =	𝑒 𝑒 𝑢 𝑖 𝑢 𝑖 𝐾 𝑖 𝐾 𝑖 𝑥 𝐴 -𝑒 𝑥 -𝑒 𝑢 𝑖 𝐾 𝑖 𝑢 𝑖 𝑥 𝐵 𝐾 𝑖 𝑥 𝐵	(𝐶 𝑡_𝐴 -𝐶 t_𝐵 ) + 𝐶 t_𝐵	(3.6)
	∆𝐶 𝑡 =	(𝑃 𝑑 -𝑤 𝑝 𝑢	𝑑𝐶 𝑝 𝑑𝑧	)	((𝑥 -𝑥 𝐵 ) + (𝑥 𝐵 -𝑥 𝐴 ) (	𝑒 𝑒 𝑢 𝑖 𝑢 𝑖 𝐾 𝑖 𝐾 𝑖 𝑥 𝐴 -𝑒 𝑥 -𝑒 𝑢 𝑖 𝐾 𝑖 𝑢 𝑖 𝑥 𝐵 𝐾 𝑖 𝑥 𝐵

  Submitted, non-definitive, version. Final published version here : https://doi.org/10.1016/j.dsr.2022.103951

Table 1 :

 1 Model parameters and output

	γn	d 231 Pap/dz	Ki	ui (m/s)	wp	Best fit	Best fit	Best fit
	(kg/m 3 )	(fg/kg/m)	(m²/s)		(m/y)	R 2 232 Th	R 2 230 Th	R 2 231 Pa
	Roy-Barman et al. (2019)
	27.865	--	2000	-0.0033	674	0.994	0.945
	(S1-S4)		±840	± 0.0014	±250
	28.094	--	2180	-0.0036	418	0.696	0.806
	(S3-S5)		±480	± 0.0008	±470
	This study
	27.865	3.7e-6	1900	-0.0030	990 ±85 0.994	0.945	0.997
	(S1-S4)	±1.9e-6	±180	±0.0003
	28.094	2.5e-4	1900	-0.0031	370	0.0696	0.806	0.922
	(S3-S5)	±2.3e-4	±120	±0.0002	±200

  Table ES1, we obtain:

						𝐹 𝑑𝑖𝑓𝑓 𝐹 𝑝𝑟𝑜𝑑	=	𝑙 × ℎ × 𝑃 × 𝑙 × ℎ × 𝐿 231 𝑃𝑎 𝑑 𝐿 × 𝐾 𝑖	=	231 𝐾 𝑖 × 𝑃𝑎 𝑑 𝑃 × 𝐿²
						𝐹 𝑎𝑑𝑣 𝐹 𝑝𝑟𝑜𝑑	=	𝑙 × ℎ × 𝑢 𝑖 × 𝑃𝑎 𝑑 231 𝑃 × 𝑙 × ℎ × 𝐿	=	𝑃𝑎 𝑑 231 𝑃 × 𝐿 × 𝑢 𝑖
					𝐹 𝑠𝑐𝑎𝑣 𝐹 𝑝𝑟𝑜𝑑	=	𝑙 × 𝐿 × 𝑤 𝑝 × 𝑃𝑎 𝑝 231 𝑃 × 𝑙 × ℎ × 𝐿	=	𝑤 𝑝 × ℎ × 𝑃 231 𝑃𝑎 𝑝
	𝐹 𝑑𝑖𝑓𝑓 𝐹 𝑝𝑟𝑜𝑑	,	𝐹 𝑎𝑑𝑣 𝐹 𝑝𝑟𝑜𝑑	𝑎𝑛𝑑	𝐹 𝑠𝑐𝑎𝑣 𝐹 𝑝𝑟𝑜𝑑	are given in Tab.

Table ES1 :

 ES1 Box-model parameters used and results

	Submitted, non-definitive, version. Final published version here : https://doi.org/10.1016/j.dsr.2022.103951
		γn = 27.865 kg/m3	γn = 28.094 kg/m3
	L (m)	800000	500000
	h (m)	1429	1205
	P (fg/m 3 /y)	25	25
	ui (m/s)	0.0030	0.0031
	Ki (m²/s)	1900	1900
	wp (m/y)	990	340
	231 Pad (fg/m 3 )	2000	2100
	231 Pap (fg/m 3 )	15	18
	𝐹 𝑎𝑑𝑣 𝐹 𝑝𝑟𝑜𝑑	946%	1642%
	𝐹 𝑑𝑖𝑓𝑓 𝐹 𝑝𝑟𝑜𝑑	749%	2013%
	𝐹 𝑠𝑐𝑎𝑣 𝐹 𝑝𝑟𝑜𝑑	41%	20%

Table ES2 :

 ES2 Dissolved concentration of 231Pa from samples of the Bonus GoodHope cruise

		Depth	θ	γ n	231 Pad	231 Pad-xs	DSi
	Station	(m)	(°C) Salinity	(kg/m 3 )	(fg/kg)	(fg/kg)	(µg/kg)

Table ES3 :

 ES3 Particulate 231Pa from Bonus GoodHope samples and the FTh/Pa measured from the particulate measurement and a linear interpolation over dissolved sample Depth (m) 231 Pap (fg/kg) 232 Thp (fg/kg) 231 Pap-xs (fg/kg) BSi (µg/kg) FTh/Pa Submitted, non-definitive, version. Final published version here : https://doi.org/10.1016/j.dsr.2022.103951

	Station S1
	30	0.006 ±0.001	7.0 ±0.1	0.005 ±0.001	18.82	2.41 ±0.55
	199	0.005 ±0.001	18.8 ±0.3	0.004 ±0.001		10.53 ±1.76
	1242	0.017 ±0.003	32.3 ±0.7	0.016 ±0.003		26.55 ±4.19
	2687	0.020 ±0.003	37.8 ±0.5	0.018 ±0.003		15.52 ±2.48
	2732	0.015 ±0.002	17.3 ±0.5	0.015 ±0.002		27.41 ±7.89
	4628	0.341 ±0.054	371 ±10	0.325 ±0.052		10.60 ±1.74
	Station S2
	20	0.017 ±0.004 4.1	±0.1	0.016 ±0.004	2.7	2.95	±0.7
	257	0.007 ±0.003 8.9	±0.1	0.007 ±0.003		13.10 ±5.99
	601	0.007 ±0.002 5.4	±0.1	0.007 ±0.002		14.61 ±4.11
	1426	0.010 ±0.004 9.1	±0.1	0.010 ±0.004		37.63 ±16.25
	1941	0.006 ±0.001 8.5	±0.1	0.006 ±0.001		25.82 ±4.16
	2859	0.014 ±0.002 8.9	±0.1	0.013 ±0.002		23.07 ±5.24
	3858	0.106 ±0.017 93.1 ±1.2	0.103 ±0.016		13.19 ±2.14
	Station S3
	40	0.003 ±0.001 7.1	±0.1	0.003 ±0.001	21.50	16.49 ±5.87
	594	0.018 ±0.002 4.41 ±0.04	0.017 ±0.002	6.72	8.452 ±1.24
	1058	0.020 ±0.003 5.1	±0.1	0.020 ±0.003	3.36	12.41 ±1.72
	1985	0.023 ±0.003 5.5	±0.1	0.023 ±0.003		12.58 ±1.80
	3025	0.027 ±0.004 8.0	±0.1	0.026 ±0.004		12.53 ±2.11
	4262	0.021 ±0.003 2.5	±0.0	0.021 ±0.003		3.12	±0.47
	Station S4
	59	0.001 ±0.000 0.076 ±0.001 0.001 ±0.000	20.16	3.83	±0.90
	169	0.019 ±0.001 1.48 ±0.02	0.019 ±0.001	24.86	2.66	±0.54
	258	0.028 ±0.004 2.42 ±0.02	0.028 ±0.004	18.82	2.23	±0.35
	743	0.030 ±0.001 2.23 ±0.02	0.030 ±0.001	8.74	3.88	±0.21
	1119	0.019 ±0.000 2.82 ±0.03	0.019 ±0.000	9.41	4.67	±0.26
	1682	0.013 ±0.002 2.23 ±0.02	0.012 ±0.002	2.02	7.89	±1.19
	2273	0.082 ±0.011 6.07 ±0.06	0.082 ±0.011	7.39	3.63	±0.52
	2469	0.115 ±0.016 6.25 ±0.06	0.115 ±0.015
	Station S5
	30	0.027 ±0.004 2.23 ±0.01	0.027 ±0.004
	139	0.023 ±0.001 2.1	±0.1	0.023 ±0.001	8.064	0.57	±0.06
	248	0.079 ±0.003 2.5	±0.1	0.079 ±0.003	15.46	0.57	±0.03
	694	0.043 ±0.007 3.6	±0.8	0.043 ±0.007	6.72	1.05	±0.20
	1784	0.031 ±0.003 4.0	±0.4	0.031 ±0.003	3.36	1.81	±0.20
	2478	0.063 ±0.002 8.5	±0.1	0.063 ±0.002	2.69	1.68	±0.10
	3172	0.089 ±0.003 11.6 ±0.2	0.089 ±0.003	2.69	1.58	±0.08
	3840	0.179 ±0.006 26.1 ±0.3	0.178 ±0.006		1.62	±0.08

Table ES4 :

 ES4 Major phases in filtered particles and Th and Pa partition coefficients for the bulk particulate matter and Pa partition coefficient estimated for pure opal.BSi from[START_REF] Fripiat | Silicon pool dynamics and biogenic silica export in the Southern Ocean inferred from Si-isotopes[END_REF], POM estimated from the POC[START_REF] Cavagna | Water column distribution and carbon isotopic signal of cholesterol, brassicasterol and particulate organic carbon in the Atlantic sector of the Southern Ocean[END_REF] and232 Th from Roy-Barman and al. 2019 * The Kd-opal( 231 Pa) reported here are only ones with the opal contribution to the total partition is significant toward the other phases of the particulate matter

	station depth	Bsi	CaCO3	232 Th	POM	SPM	% Opal Kd-bulk(Pa)	Kd-bulk(Th)	Kd-opal(Pa)*
		(m)	(µg/kg)	(µg/kg)	(pg/kg) (µg/kg) (µg/kg)	(g/g) ×10 6 (g/g) ×10 6 (g/g) ×10 6 (g/g)
	1	25	18,82	33	6.96	26.7	79.2	24 %	0.378	0,536	-
	2	70	2,69	19	6.49	45.6	67.9	4 %	0.741	1,803	1.62
	3	41	21,50	55,2	7.06	75.4	152.8	14 %	0.062	0,757	-
	3	600	6,72	5,7	4.41	7.4	20.2	33 %	0.719	5,848	0.59
	3	1068	3,36	4,3	5.07	4.4	12.6	27 %	0.826	9,920	0.81
	4	60	20,16	19	0.07	40	79.2	25 %	0.054	0,212	-
	4	170	24,86	2,4	1.39	12.0	39.4	63 %	0.975	2,692	1.16
	4	260	18,82	6,3	2.37	8.4	33.8	56 %	1.225	2,721	1.60
	4	749	8,74	2,5	2.21	5.6	17.0	51 %	0.851	3,274	0.96
	4	1128	9,41	2,5	2.76	6.0	18.1	52 %	0.474	2,243	0.23
	4	1695	2,02	17	2.14	5.8	25.0	8 %	0.214	1,726	-
	4	2305	7,39	12	6.06	5.7	25.7	29 %	1.125	4,029	1.80
	5	140	8,06	6,4	2.09	20.2	34.9	23 %	1.578	0,881	4.55
	5	250	15,46	2,1	2.42	6.5	24.3	64 %	2.279	1,315	3.17
	5	700	6,72	3,1	3.48	9.5	19.7	34 %	1.153	1,199	2.00
	5	1800	3,36	12,1	3.93	1.8	17.7	19 %	0.737	1,324	0.04
	5	2500	2,69	58	8.31	3.0	64.5	4 %	0.460	0,768	-
	5	3200	2,69	7	11.03	5.1	15.9	17 %	1.962	3,151	7.15

Thm are the measured isotopes[START_REF] Roy-Barman | Thorium isotopes in the Southeast Atlantic Ocean: Tracking scavenging during water mass mixing along neutral density surfaces[END_REF]. λ230, λ231, λ235 and λ238 are the decay constant of 230 Th, 231 Pa, 235 U and 238 U respectively. M231 and M230 are the molar masses Submitted, non-definitive, version. Final published version here : https://doi.org/10.1016/j.dsr.2022.103951

Submitted, non-definitive, version. Final published version here : https://doi.org/10.1016/j.dsr.2022.103951

Acknowledgements

Authors would like to acknowledge all the on-board crew and sampling team of BGH cruise. This work also supported by the French National program LEFE (Les Enveloppes Fluides et l'Environnement) through support to the AbAc project.