Claypool continued: Extending the isotopic record of sedimentary sulfate

Chemical Geology 513 (2019) 200–225

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Invited review article

Claypool continued: Extending the isotopic record of sedimentary sulfate a,b,c,⁎

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Peter W. Crockford , Marcus Kunzmann , Andrey Bekker , Justin Hayles , Huiming Bao , Galen P. Halversona,i, Yongbo Pengg, Thi H. Buia, Grant M. Coxj, Timothy M. Gibsona, Sarah Wörndlea, Robert Rainbirdk, Aivo Leplandl, Nicholas L. Swanson-Hysellm, Sharad Mastern, Bulusu Sreenivaso, Anton Kuznetsovp, Valery Krupenikq, Boswell A. Winga,r

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Department of Earth and Planetary Sciences, McGill University, Montreal, QC, Canada Department of Earth and Planetary Sciences, Weizmann Institute of Science, 76100 Rehovot, Israel c Department of Geoscience, Princeton University, Princeton, NJ 08544, USA d CSIRO Mineral Resources, Australian Resources Research Centre, Kensington, WA 6151, Australia e University of California at Riverside, Riverside, CA 92521, USA f Department of Earth Science, MS-126, Rice University, P.O. Box 1892, Houston, TX 77251, USA g Department of Geology and Geophysics, Louisiana State University, Baton Rouge, LA 70803, USA h School of Earth & Space Sciences, Peking University, Beijing 100871, China i Earth Dynamics Research Group, ARC Centre of Excellence for Core to Crust Fluid Systems (CCFS), The Institute for Geoscience Research (TIGeR), School of Earth and Planetary Sciences, Curtin University, GPO Box U1987, WA 6845, Australia j Department of Earth Sciences, The University of Adelaide, Adelaide, SA 5005, Australia k Natural Resources Canada, Geological Survey Canada, 601 Booth St, Ottawa, ON K1A 0E8, Canada l Geological Survey of Norway, Trondheim, Norway m Department of Earth and Planetary Science, University of California, Berkeley, CA 94720, USA n EGRI, School of Geosciences, University of the Witwatersrand, P. Bag 3, WITS 2050 Johannesburg, South Africa o CSIR-National Geophysical Research Institute, National Geophysical Research Institute, Hyderabad, India p Institute of Precambrian Geology and Geochronology (IPGG), Russian Academy of Sciences, Saint Petersburg 199034, Russia q Sector of Precambrian Geology, A.P. Karpinsky Russian Geological Research Institute, 74 Sredny Prospect, 199106 St. Petersburg, Russia r Department of Geological Sciences, University of Colorado Boulder, UCB 399, Boulder, CO 80309–0399, USA b

ARTICLE INFO

ABSTRACT

Editor: Michael E. Böttcher

The Proterozoic Eon spans Earth's middle age during which many important transitions occurred. These transitions include the oxygenation of the atmosphere, emergence of eukaryotic organisms and growth of continents. Since the sulfur and oxygen cycles are intricately linked to most surface biogeochemical processes, these transitions should be recorded in changes to the isotopic composition of marine and terrestrial sulfate minerals. Here we present oxygen (∆17O, δ18O) and sulfur (∆33S, δ34S) isotope records of Proterozoic sulfate from currently available data together with new measurements of 313 samples from 33 different formations bearing Earth's earliest unambiguous evaporites at 2.4 Ga through to Ediacaran aged deposits. This record depicts distinct intervals with respect to the expression of sulfate isotopes that are not completely captured by established intervals in the geologic timescale. The most salient pattern is the muted ∆17O signatures across the GOE, late Proterozoic and Ediacaran with values that are only slightly more negative than modern marine sulfate, contrasting with highly negative values across the mid-Proterozoic and Cryogenian. We combine these results with estimates of atmospheric composition to produce a gross primary production (GPP) curve for the Proterozoic. Through these results we argue that changes in GPP across Earth history likely help account for many of the changes in the Proterozoic Earth surface environment such as rising atmospheric oxygen, large fluctuations in the size of the marine sulfate reservoir and variations in the isotopic composition of sedimentary sulfate.

Keywords: Gross primary production Marine sulfate Proterozoic Precambrian Sulfate Triple oxygen Oxygen isotopes Sulfur isotopes Multiple sulfur Isotope geochemistry Evaporite Gypsum Barite Atmospheric oxygen Primary production Primary productivity Evolution of life Biosphere

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Corresponding author at: Department of Earth and Planetary Sciences, Weizmann Institute of Science, 76100 Rehovot, Israel E-mail address: [email protected] (P.W. Crockford).

https://doi.org/10.1016/j.chemgeo.2019.02.030 Received 25 May 2018; Received in revised form 28 January 2019; Accepted 17 February 2019 Available online 07 March 2019 0009-2541/ © 2019 Elsevier B.V. All rights reserved.

P.W. Crockford, et al.

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1. Introduction

2003; Johnston, 2011; Kunzmann et al., 2017). Sulfate (SO42−) is the second most abundant anion in modern marine environments, and although its abundance was much lower through most of Earth's history (Kah et al., 2004; Johnston et al., 2008; Canfield and Farquhar, 2009; Bekker and Holland, 2012; Crowe et al., 2014; Luo et al., 2015), it must have played a significant role in ancient biogeochemical cycles (Canfield and Raiswell, 1999). Sulfate is also the only oxy-anion known to be capable of preserving ancient atmospheric oxygen and sulfur isotope ratios, providing a window into atmospheric chemistry as well as the productivity of the ancient biosphere (Farquhar et al., 2000; Bao et al., 2008; Bao, 2015; Crockford et al., 2018). Claypool et al. (1980) presented the first comprehensive survey of the

The Earth System has dramatically evolved across its 4.6 Ga history (Fig. 1) experiencing massive changes to the surface environment, including the oxygenation of the atmosphere (Farquhar et al., 2000; Holland, 2006; Bekker and Holland, 2012; Lyons et al., 2014), snowball glaciations (Hoffman et al., 1998; Kirschvink et al., 2000; Bekker, 2014a), and the evolution of the biosphere (El Albani et al., 2010; Brocks et al., 2017; Knoll and Nowak, 2017; Gibson et al., 2017). It is reasonable to assume that these events significantly impacted the cycles of oxygen, sulfur, iron and carbon, and hence the sedimentary sulfate isotope record (Rouxel et al., 2005; Farquhar and Wing,

Fig. 1. Schematic evolution of the Earth System over geologic time. At the top we track changes in solar output relative to present levels calculated from Gough (1981). Below we track proposed trajectories of atmospheric CO2 and O2 levels presented relative to Present Atmospheric Levels (PAL; 280 ppm CO2; 200,000 ppm O2). The CO2 field is taken from the 1D model of von Paris et al. (2008) from 4.6 to 0.6 Ga, and using estimates from Franks et al. (2014) and Berner (2006). The O2 field is based on combined proxy data compiled by Lyons et al., 2014 with average estimates in light blue, and a broader range presented in purple/dark blue. Overlaying atmospheric estimates are panglacial intervals (so called Snowball Earth events) in blue lines (Hoffman et al., 1998; Kirschvink et al., 2000). The biosphere panel depicts changes in maximum body sizes of organisms over Earth history, as summarized by Payne et al. (2011), and overlain with the diversity of the biosphere separated into prokaryote, eukaryote and animals with dashed bars representing uncertainty in origin. In the hydrosphere panel we plot marine redox conditions (Hardisty et al., 2017; Anbar et al., 2007; Planavsky et al., 2011) together with a predicted evolution of seawater pH values (Halevy and Bachan, 2017). Finally, in the geosphere panel we plot the distribution of passive margins through time (dark brown; Bradley, 2011), supercontinents (and cratonic amalgamations e.g. Sclavia/Superia) through Earth history, and crustal growth curves from Jacobsen (1988) (red), Ying et al. (2011) (green) and Taylor and McLennan (1985) (blue). 201

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major isotopes of sulfur (δ34S) and oxygen (δ18O) within sulfate over the last one billion years of Earth history. This record provided a foundation to evaluate secular variations in seawater sulfate concentrations and to constrain how the cycles of sulfur and oxygen operated over this interval. Since this pioneering work, much additional effort has gone into filling in the δ18O and δ34S records and extending them to earlier times (e.g. Strauss, 1993). The results have highlighted both links and disconnects between these two isotopic records (Utrilla et al., 1992; Strauss, 1999; Kampschulte and Strauss, 2004; Bottrell and Newton, 2006; Turchyn et al., 2009; Wu et al., 2014). Advances in analytical capabilities have added new dimensions to interrogating these records through the ability to measure the minor isotopes of sulfur (33S, and 36S; Farquhar et al., 2000; Farquhar and Wing, 2003; Johnston, 2011) and oxygen (17O; Thiemens and Heidenreich, 1983; Luz et al., 1999; Thiemens, 2006; Bao, 2006). The new minor isotope datasets generated from sulfate-bearing sedimentary archives such as barite, gypsum and carbonate-associated sulfate (Bao et al., 2008, 2009; Crockford et al., 2018) as well as those bearing sulfide (Berner, 1984; Fischer et al., 2014; Scott et al., 2014; Kunzmann et al., 2017) and organic-bound sulfur (Bontognali et al., 2012; Raven et al., 2016, 2018) are rapidly improving our understanding of oxygen and sulfur cycling in the surface environment along with other processes that these isotopic records track. Nevertheless, because these techniques are more analytically demanding than traditional methods, the existing record is more fragmentary. This shortfall in data is particularly

notable for the Proterozoic where large gaps exist particularly with respect to oxygen (δ18O, ∆17O) and multiple sulfur isotopic data (∆33S, ∆36S). One notable example is the potential for mass-independent oxygen isotope anomalies to provide quantitative constraints on ancient gross primary productivity (GPP; Crockford et al., 2018), a parameter that likely underlies many of the geochemical trends observed across the sedimentary record (Des Marais et al., 1992; Brasier and Lindsay, 1998; Partin et al., 2013; Lyons et al., 2014; Planavsky et al., 2018). Here we briefly review the progress made to date in understanding Earth's ancient sulfur and oxygen cycles viewed through isotopic records of sedimentary sulfate as preserved in evaporite minerals and carbonate associated sulfate. A suite of 313 samples from 33 sedimentary formations spanning the Proterozoic were analyzed, and these data allow us to extend δ34S and δ18O age curves of sulfate through to the earliest Paleoproterozoic. These results are presented alongside new minor isotope (∆17O, and ∆33S) measurements. We then utilize these compilations to explore secular variations and links in all isotopic systems, highlighting potential causal mechanisms that may have driven revealed trends. Finally, we place a large degree of emphasis on the ∆17O record to generate a GPP curve spanning from the earliest Proterozoic to the modern and explore the implications of these results in the context of the emerging model for the evolution of the surface Earth (Bekker and Holland, 2012; Bekker, 2014a; Lyons et al., 2014; Payne et al., 2011; Planavsky et al., 2011; Sperling et al., 2015).

Box 1 Isotopic notation (Bao et al., 2016; Cao and Liu, 2011; Hayles et al., 2017; Luo et al., 2010; Matsuhisa et al., 1978; Pack and Herwartz, 2014).

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2. Fidelity of sulfate-bearing archives

concentrations of sulfate are prone to contamination through oxidation of trace pyrite within samples (Marenco et al., 2008). Contamination issues also extend to minor isotopes where it has been shown that surface weathering can shift CAS ∆17O isotopic compositions (Peng et al., 2014). Nonetheless, there are examples of comparable isotopic results generated from co-deposited CAS and evaporites (Kah et al., 2004; Gill et al., 2008), encouraging the continued, but cautionary use of CAS as an archive of the isotopic composition of seawater sulfate (e.g., Luo et al., 2015; Fike et al., 2006; Tostevin et al., 2017).

While a great deal of effort has been applied to identifying primary isotopic signatures of marine, microbial and atmospheric reservoirs (e.g., Canfield, 2001; Johnston et al., 2005a; Sim et al., 2011; Johnston, 2011; Bradley et al., 2016), reliably screening such measurements within sulfate-bearing archives (gypsum, anhydrite, barite, carbonate associated sulfate) from compromised records that have been subjected to post-depositional alteration remains a challenge (Fike et al., 2015). Given that the expression of primary isotopic signals within sulfate reflects environmental conditions that include atmospheric chemistry (e.g., Farquhar et al., 2000), marine sulfate levels (e.g., Habicht et al., 2002), and the nature and rate of microbial sulfur cycling (e.g., Leavitt et al., 2013; Wing and Halevy, 2014), variations in these parameters over Earth history result in a wide range of plausible isotopic signatures that can be interpreted to be primary in origin. The main focus of this work is sulfate evaporite minerals, but this dataset is supplemented by a subset of measurements on carbonate-associated sulfate.

2.3. Barite Sedimentary barite has been widely relied upon as an archive of both the Archean biosphere (Shen et al., 2001; Ueno et al., 2006; Horner et al., 2017) and the Phanerozoic sulfur cycle (Paytan et al., 1998, 2004; Turchyn and Schrag, 2006). However, its utility in providing insight into the Proterozoic surface environment is much less explored than other sulfate archives (e.g., Strauss and Schieber, 1990; Deb et al., 1991; Clark et al., 2004). Unlike the modern environment where barite deposits form in cold seep, diagenetic, hydrothermal, and pelagic environments (Jewell, 2000; Horner et al., 2017; Gong et al., 2018), most documented occurrences of Proterozoic barites are stratiform deposits of hydrothermal or cold seep origin (e.g., Strauss and Schieber, 1990; Deb et al., 1991; Torres et al., 2003; Clark et al., 2004; Huston and Logan, 2004; Lyons et al., 2009) that have been subjected to substantial metamorphism. Consequently, inferring modes of deposition based on observed diagnostic isotopic signatures from different modern environments (e.g. Sakai, 1971; Antler et al., 2015) is tenuous when applied to these Proterozoic aged deposits.

2.1. Evaporites Sulfate evaporites have been a feature of the sedimentary record since the Great Oxidation Event (GOE; Holland, 2002). Since basin restriction and evaporative conditions are a requirement for the precipitation of sulfate salts, a challenge in utilizing this archive in reconstructing ancient seawater sulfate composition is to decipher global from local (restricted basin) or post-depositional signatures (Claypool et al., 1980; Van Stempvoort and Krouse, 1994; Lu et al., 2001). Local effects include microbial sulfur cycling and influences from local inputs of sulfate from weathered evaporites (Lu et al., 2001) or hydrothermal brines (Longinelli and Craig, 1967). Further isotopic modification can occur due to reservoir effects coupled to fractionations of up to +1.6‰ in δ34S and 3.7‰ in δ18O between sulfate-bearing fluids and sulfate precipitates (Thode and Monster, 1965; Lloyd, 1968; Raab and Spiro, 1991). Interpreting evaporite records can also be complicated by postdepositional processes that may influence isotopic values (e.g. Fike et al., 2015) such as metamorphism that tends to shift δ18O to more positive values, and dilute primary signatures of original sulfate (Alonso-Azcárate et al., 2006). Additionally, given the high solubility of gypsum (Halevy et al., 2012; Canfield, 2013), it is unclear to what degree the temporal distribution of preserved evaporite deposits reflects secular changes in sulfate and calcium concentrations in the ocean, changes in the global configuration of sedimentary basins, or inconsistent preservation of such deposits (Mackenzie and Garrels, 1971). Despite these considerations, evaporite minerals remain among the best archives in the sedimentary record of ancient seawater (Claypool et al., 1980) or atmospheric chemistry (Crockford et al., 2018), and thus are an important archive of the ancient sulfur and oxygen cycles.

3. Isotopes of sedimentary sulfate 3.1. ∆17O Upon reaching sufficient levels of atmospheric oxygen to establish an ozone (O3) layer (pO2 = 10−3 pre-anthropogenic levels, hereafter PAL; Kasting and Donahue, 1980; Segura et al., 2003), O2 is imprinted with a mass-independent fractionation signature (O-MIF: ∆17O ≠ 0) imparted through the formation and destruction of O3 (Fig. 2). During photolysis, O3 dissociates into a single oxygen atom and one O2 molecule. Symmetry effects during recombination of O3 as well as reactions with other atmospheric species that temporarily sequester heavy isotopes, has a net effect on the O2 molecules retaining 17Oe16O bonds at a smaller proportionate amount relative to what is predicted from massdependent processes (Thiemens and Heidenreich, 1983; Heidenreich III and Thiemens, 1986; Thiemens, 2006; Fig. 2). Since initial experiments exploring the photochemical dissociation of ozone, it has been observed that the magnitude of the ∆17O signal in atmospheric oxygen is also dependent upon reactions involving the spalled-off oxygen atom and CO2, as well as other stratospheric species that sequester mass independent enrichment of heavy oxygen isotopes (17O, 18O) (Wen and Thiemens, 1993; Blunier et al., 2002; Bao et al., 2008). The result is that the magnitude of ∆17O signals in atmospheric O2 is proportional to CO2 concentrations (pCO2) (Gamo et al., 1989; Wen and Thiemens, 1993; Yung et al., 1991, 1997; Fig. 2). The flux of O-MIF from the stratosphere is counteracted in the troposphere by photosynthetically produced O2 from the biosphere that bears a ∆17O value of seawater (0‰; Luz et al., 1999). Although dependent upon the proportion of O2 that is both directly produced through autotrophic carbon fixation as well as the proportion of oxygen that interacts with the atmospheric O2 reservoir before respiration, this biospheric oxygen flux has been shown to approximate gross primary production (hereafter GPP) and for the purposes of the present study we remain consistent with this convention (e.g., Luz et al., 1999; Blunier et al., 2002; Crockford et al., 2018). Therefore, the ∆17O signature of atmospheric oxygen represents a balance between the amount of CO2

2.2. Carbonate associated sulfate (CAS) Due to the sparse distribution of evaporites through much of the Proterozoic, as well as the difficulty in dating such deposits, many efforts have focused on sulfate bound within carbonates (e.g., Hurtgen et al., 2002; Jones and Fike, 2013; Guo et al., 2009; Luo et al., 2015). A challenge with CAS is the difficulty in disentangling the influences of pore-water processes and sedimentation rates versus original seawater sulfate composition on preserved isotopic values (Rennie and Turchyn, 2014; Fike et al., 2015). Recent work highlights that δ18O values are less resilient to post-depositional alteration than coeval δ34S values (Gill et al., 2008; Turchyn et al., 2009; Rennie and Turchyn, 2014; Owens et al., 2013; Gomes et al., 2016; Fichtner et al., 2017). Additional challenges can arise under low seawater sulfate conditions where local processes can exert a large influence on isotopic values. Furthermore, low CAS concentrations have led to analytical challenges, as small 203

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+∆17O OOQ

Q = 17,18O

Q + CO2 CO2Q

-∆17O O2 + Q

+∆17O

Q + CO2

seawater (0 ‰)

Stratospheric Anomaly

Pyrite Oxidation (fO2 = 8-15%)

Tropospheric Isotopic Composition

∆17O

Photosynthesis (GPP) Evaporite Weathering/ Deposition

Fig. 2. A simplified schematic of the controls on the ∆17O composition of sulfate atmospheric O2 for the Earth surface environment. In the upper portion of the figure, the photochemical network of reactions is depicted where yellow Q's represent 17 O and 18O and it is shown how O2 can leave the stratosphere with a depletion in this isotopically heavy oxygen (Thiemens and Heidenreich, 1983; Wen and Thiemens, 1993; Yung et al., 1991, 1997; Luz et al., 1999). Arrows change colour from blue to red to indicate processes that impart the troposphere or sulfate with negative ∆17O values (blue) and those that push ∆17O values toward zero (red), with process that behave neutrally in yellow such as pyrite oxidation (Balci et al., 2007).

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Microbial S Cycling

available to sequester the positive ∆17O anomaly, the rate of gross oxygen production from the biosphere (i.e. GPP), and the size of the O2 reservoir (pO2) where these fluxes are mixed (Cao and Bao, 2013; Fig. 2). This combination of factors results in modern tropospheric O2 being mass independently depleted in 17O with a ∆17O value of −0.469‰ (Pack et al., 2017; Barkan and Luz, 2011). The most straightforward example of the CO2-∆17O relationship is observed in the ice core record for the past 60,000 years where increases in CO2 from the last glacial period to the present are paralleled by decreasing ∆17O values (Blunier et al., 2002). Given that pristine atmospheric oxygen archives do not extend beyond the ice-core record (< 1 Myrs; Barnola et al., 1987; Petit et al., 1999; Higgins et al., 2015; Stolper et al., 2016), other oxygen-bearing archives that can preserve a portion of the atmospheric signal are required to explore earlier times in Earth history. Sulfate is one such archive, and it benefits from greater immunity to isotopic exchange than many other oxy-anions (Hall and Alexander, 1940; Gamsjager and Murmann, 1983; Bao, 2015). A predictable fraction of atmospheric oxygen is incorporated into product sulfate (8–30%) during the oxidation of sulfide minerals on the continents and therefore a predictable proportion of the atmospheric O2 ∆17O value (Balci et al., 2007; Kohl and Bao, 2011; Fig. 2). In environments where this ∆17O signal is not erased by sulfur cycling, it might be transferred to the sedimentary record and archived for over a billion years (e.g. Crockford et al., 2018). Importantly, all processes following sulfide oxidation will remove or dilute anomalous ∆17O values in sulfate, meaning that ∆17O values of sulfate minerals provide a conservative estimate of original ∆17O values of atmospheric O2. Despite this potential utility of ∆17O signatures in probing ancient pCO2 and GPP levels, which are inferred to have varied significantly through Earth history (Fig. 1), the current ∆17O record of sulfate hardly extends beyond the Cryogenian (717–635 Ma; Bao et al., 2008; Bao, 2015) with only few earlier data sets (Blättler et al., 2018; Crockford et al., 2018).

3.2. δ18O The δ18O composition of modern marine sulfate (≈8.6‰; Brand et al., 2009; Lloyd, 1968; Johnston et al., 2014) sits between the two large biologically accessible reservoirs: atmospheric O2 at 22.9‰ (Nier, 1950) and seawater at 0‰ (Fig. 3). If allowed to reach equilibrium with seawater at 25 °C, sulfate would be enriched in 18O by ~23‰ relative to seawater (Zeebe, 2010); however, the slow kinetics of isotope exchange at low temperatures and moderate pH, allows for various processes to modify its isotopic composition away from this equilibrium value (Zak et al., 1980; Turchyn and Schrag, 2006). For example, varying the degree of dissimilatory sulfate reduction might be manifested in changes in sulfate δ18O values where intermediates such as thiosulfate and sulfite readily exchange oxygen with water, and thus drive the sulfate pool to be isotopically enriched (Mizutani and Rafter, 1973; Fritz et al., 1989; Wortmann et al., 2007; Antler et al., 2017). However, many details are not understood about these processes (Antler and Pellerin, 2018). An example of an abiological mechanism known to shift the δ18O of seawater sulfate to more positive values is the dissolution of isotopically enriched ancient evaporite deposits (Tostevin et al., 2014; Fig. 3). This mechanism, while operational for modern environments, may not be as relevant for earlier times in Earth history where evaporites were likely deposited under different pH, pCO2 and marine sulfate concentration conditions (Fig. 1). Finally, the production of dimethyl sulfide (DMS) through the degradation of dimethylsulfonium propionate in marine algae leads to the rapid exchange of isotopes between SO2 and other oxy-anions that can enrich product sulfate by up to 20‰ in δ18O (Kumar et al., 2002; Holt et al., 1983). This flux has been estimated to be up to one third of the flux of modern riverine sulfate (Turchyn and Schrag, 2006), meaning oxygenisotopic records of sulfate might reflect the emergence and expansion of this and other metabolisms that oxidize sulfur (Figs. 1 and 3). Direct sulfide oxidation on the continents and reoxidation in

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+

Fig. 3. Schematic of the δ18O system and its interpretation within sulfate. We outline the dominant controls on marine sulfate following Turchyn and Schrag (2006). Text in red indicates processes that will drive δ18O values more positive, text in blue indicates processes that will drive δ18O values more negative, and yellow indicates a process that can either drive values more positive or negative or does not impart any significant fractionation.

Isotopic exchange with water vapor

δ18O DMSO Production <650 Ma

Sulfide Weathering Evaporite Weathering/ Deposition

-

S-Reoxidation

sediments are thought to be the primary processes driving the δ18O sulfate composition to values typically lower than initial marine sulfate values (Van Stempvoort and Krouse, 1994). However, it may also be possible for biologically mediated sulfide oxidation to drive δ18O values more positive in some environments (Böttcher et al., 2001). Importantly the influence of reoxidative sulfur cycling along with other processes that impact the δ18O of sedimentary sulfate have likely varied significantly through Earth history (Canfield and Farquhar, 2009; Tarhan et al., 2015; Kunzmann et al., 2017; Fig. 1). Therefore, both abiological and biologically mediated isotopic exchange between these reservoirs governs the δ18O composition of sulfate in modern environments.

Microbial S Cycling

laboratory studies that monitor dissimilatory sulfate reduction, it has been argued that this metabolism has been a dominant control on the sedimentary sulfur isotope record for much of Earth history (Shen et al., 2001; Butler et al., 2004; Johnston, 2011; Leavitt et al., 2013; Fig. 4). Another metabolic pathway that imparts a δ34S signature to marine sulfate involves inorganic fermentation without phosphorylation. This metabolism disproportionates intermediate sulfur species (e.g., elemental sulfur, thiosulfate, and sulfite) to produce isotopically distinct H2S and sulfate (Bak and Cypionka, 1987; Jørgensen, 1990; Thamdrup et al., 1993; Canfield et al., 1998; Finster, 2008) and it is suggested that this metabolism only rose to prominence in the early Ediacaran (Kunzmann et al., 2017). In the absence of significant sulfur disproportionation, Proterozoic sulfur isotopic records are likely to be controlled by the intensity of dissimilatory sulfate reduction that over time can lead to a progressively isotopically heavier marine sulfate reservoir due to preferential utilization of light isotopes (Thode et al., 1951; Harrison and Thode, 1958; Fig. 4). Therefore, the types of sulfur metabolisms, extent of their ability to cycle sulfur, and relative strength of abiological processes (such as weathering and hydrothermal reactions) control the δ34S composition of the marine sulfate reservoir (Kah et al., 2004; Bottrell and Newton, 2006; Gomes and Hurtgen, 2013; Tostevin et al., 2014; Fig. 4).

3.3. δ34S Sulfate primarily enters the ocean via riverine input, which has a δ34S composition that reflects provenance in the watershed (Killingsworth et al., 2018; Burke et al., 2018). This composition can be quite variable depending on the dominant lithology being weathered, (i.e. sulfate evaporites, sulfidic shales, or crustal sulfides; Fig. 4). In the modern environment, and likely during the Proterozoic, the long term removal of sulfur from the ocean is achieved through pyrite precipitation and burial (≈10–99%) and to a lesser degree as sulfate minerals, since sulfate-evaporite deposits are prone to rapid recycling (Canfield, 2004; Halevy et al., 2012; Canfield, 2013; Tostevin et al., 2014; Fig. 4). Additional fluxes of sulfate into and out of the marine reservoir include biologically and abiologically mediated sulfide oxidation (source), volcanic inputs (source) and hydrothermal alteration of oceanic crust (sink; Wolery and Sleep, 1976; Alt, 1995; Fig. 4). Pyrite burial is controlled by iron availability in sediments where sulfate is effectively respired to produce H2S/HS− through dissimilatory sulfate reduction, and then reacts with ferrous iron to produce iron sulfide minerals (in most cases pyrite; Jørgensen, 1982). Although transport processes and reservoir effects can exert a large influence on preserved records, since the isotopic difference between co-existing pyrite and sulfate minerals is similar to sulfur isotope fractionations commonly expressed during

3.4. ∆33S The near disappearance of mass-independent sulfur isotope anomalies from the geologic record in the earliest Paleoproterozoic (ca. 2.4 Ga) has provided critical constraints on the evolution of atmospheric oxygen across the GOE (Farquhar et al., 2000; Pavlov and Kasting, 2002; Bekker et al., 2004). Before the GOE, there was an absence of an ozone layer (Segura et al., 2003) due to pO2 levels below 10−5 PAL (Pavlov and Kasting, 2002). This change in the atmosphere allowed for UV radiation to permeate the lower atmosphere and drive photochemistry that imparted a large mass-independent fractionation into sulfur species (S-MIF). Unlike the post-GOE world, atmospherically 205

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derived S-MIF signatures could survive sulfur cycling throughout the surface environment and thus are expressed in the sedimentary record (Halevy, 2013; Fig. 5). Beyond a litmus test for the presence of atmospheric oxygen, multiple sulfur isotopes within the mass-dependent range of fractionation has proven effective in identifying different metabolisms and processes that obey different mass laws (33λ or 33θ; cf. Johnston, 2011) that, in turn, govern their diagnostic fractionations in ∆33S-δ34S space (Farquhar et al., 2003; Johnston et al., 2005a; Sim et al., 2011; Johnston et al., 2011; Zerkle et al., 2009; Kunzmann et al., 2017). Differences in mass laws arise because different metabolisms involve variable degrees of isotopic exchange between sulfur-bearing species during respective metabolic processes, which are potentially a result of different capacities for kinetic isotope effects to modify isotopic values (Wing and Halevy, 2014). For example, sulfur disproportionators have been found to have a higher affinity for taking up 33S versus 34S during metabolic processes compared to dissimilatory sulfate reducers, resulting in a slightly larger (33λ > 0.515) mass law (Johnston et al., 2005a). Beyond metabolic processes, the use of minor sulfur isotopes can also aid in a more accurate assessment of processes operating within sediments (e.g. diffusion; Pellerin et al., 2015). However, the Proterozoic Eon is far from replete with ∆33S data for sulfates, leaving models based on the ∆33S values of sulfides reliant upon assumptions about initial marine sulfate isotopic values (Scott et al., 2014; Kunzmann et al., 2017). 4. Methods

Marenco et al., 2008), calculated modal pyrite abundances within a subset of analyzed evaporite samples of < 0.3% pyrite‑sulfur indicates any such isotopic contamination to be minor. For oxygen isotope measurements (∆17O and δ18O), samples (both drilled evaporites and precipitated barites from CAS extractions) were taken through a series of dissolution and precipitation steps (Bao, 2006) in order to remove all non-sulfate oxygen-bearing contaminants such as nitrate. Samples were first dissolved into a 0.05 N Diethylenetriaminepentaacic acid (DTPA) – 1.0 M sodium hydroxide (NaOH) solution. After the dissolution step, samples were passed through 0.2 μm filters to remove silicates or non-soluble residues in samples. After filtering, samples were precipitated as barite by increasing saturation and lowering pH, through addition of double-distilled 6 N HCl at 80 °C followed by BaCl2, thus preventing witherite (BaCO3) formation. This full procedure was then repeated and final products were dried in an oven at 80 °C. To prepare samples for sulfur isotope analysis samples (both barites from CAS extractions and evaporite minerals) were boiled at 100 °C in Thode solution (HI, H3PO2 and HCl; Thode et al., 1961; Pepkowitz and Shirley, 1951). This procedure converts sulfate to hydrogen sulfide (H2S), which is then transported through a chilled column in a nitrogen gas stream, bubbled through deionized water, and converted in a zinc acetate trap to zinc sulfide (ZnS). Solutions containing ZnS were then reacted with 0.2 M silver nitrate (AgNO3) to precipitate Ag2S. Samples containing Ag2S precipitates were collected on a 0.45 μm membrane and dried in an oven at 80 °C. Once dried, ~3 mg of Ag2S was placed into a cleaned aluminum foil for multiple S isotope analyses.

4.1. Samples

4.3. Isotope analyses

For this study we analyzed 313 samples from 33 different formations on all continents but South America and Antarctica (Figs. 6 and 7; Table 1). We rely on the most current literature estimates of ages for formations and summarize these along with sample locations and formation names in Fig. 6 and Table 1, respectively. This sample suite covers the oldest known sulfate evaporites from North America (2.32 Ga Gordon Lake Formation) and from South Africa (2.4 Ga Duitschland Formation) to Ediacaran-aged samples from southern Iran and Siberia. The majority of samples were deposited as sulfate evaporite minerals (n = 305) with only a few exceptions where carbonateassociated sulfate (CAS) was measured (n = 8). The data generated for this study are then plotted with existing data in new compilations of the Proterozoic sulfate isotope record (Figs. 7 and 8).

All oxygen isotope measurements were performed at the Louisiana State University OASIC laboratory. Oxygen was generated from sulfates to measure ∆17O values using a laser fluorination system. Approximately 10 mg of powder for each sample was loaded onto a stainless steel (SS316) plate and placed into a SS316 chamber capped with a BaF2 window. The chamber was then exposed to a bromine pentafluoride (BrF5) atmosphere of > 100 mbar for 3 min, evacuated, and injected again with BrF5 at 20 mbar for 12 h. The sample chamber was then evacuated and filled again with a fresh injection of BrF5. Samples were heated with a CO2 laser, which liberated oxygen from sulfate. Oxygen gas was purified by passing over five cold traps at −196 °C that removed any condensable gases. Purified oxygen was then collected onto 5A mol-sieve immersed in a –196 °C cold trap over a seven-minute interval. Samples were then analyzed in dual inlet mode on a Thermo MAT-253. Analyses were conducted over 3 acquisitions consisting of 8 standard-sample brackets. Error on the total analytical procedure is estimated based on repeat measurements of in-house BaSO4 calibrated against UWG-2 (taken to have δ18O = +5.80‰ and δ17O =Δ 3.016‰ on the VSMOW scale, resulting in ~0.03‰ for ∆17O with a maximum 1σ error of 0.05‰ on individual analyses). Measurements of δ18O values were performed on the same samples that had been through dissolution-precipitation steps outlined above. Between ~180 and 220 μg of sample was weighed out and wrapped in silver foil before loading samples into a thermal conversion elemental analyzer (TC/EA) coupled to the same Thermo MAT-253 isotope ratio mass spectrometer set in continuous flow mode. Samples were analyzed in duplicate, and the total error on analyses is estimated to be < 0.5‰ based on repeated measurements of laboratory standards (USGS34 δ18O = −27.80‰; IAEA-NO3 δ18O = 25.6‰; USGS35 δ18 O = +56.80‰) on the VSMOW scale. For sulfur isotope analyses, 3 mg samples of Ag2S within aluminum foil packets were placed into a nickel bomb under a fluorine gas (F2) atmosphere at 250 °C and allowed to react for 12 h. Within nickel bombs F2 reacted with Ag2S to produce sulfur hexafluoride gas (SF6).

4.2. Sample preparation Sulfate evaporites were micro-drilled from thin beds, veins and bladed crystals, or hand crushed using an agate mortar and pestle in the case of thickly bedded deposits and nodules. For CAS, 50 to 500 g of carbonate were crushed in a steel ring mill and resulting powders were then placed into a 1 L Pyrex Erlenmeyer flask. Powdered carbonates were first placed in a 5% sodium chloride (NaCl) solution for 12 h and then rinsed multiple times with deionized water to remove any nonCAS sulfate. Next, samples were placed in a weak (5%) hydrogen peroxide (H2O2) solution for 12 h in order to oxidize pyrite and subsequently remove it from solution as sulfate. Samples were then dissolved into a 4 N hydrochloric acid (HCl) - 5% tin chloride (SnCl2) solution for a 12-hour period. Samples were then decanted and filtered through a 0.45 μm filter and mixed with a concentrated barium chloride (BaCl2) solution to precipitate liberated sulfate as barite over a 72-hour period. Finally, barite precipitates were collected onto a 0.2 μm filter, lightly rinsed with 4 N HCl, and dried in an oven at 80 °C. Although HCl and acetic acid have been shown to oxidize sulfides, and thus potentially affect both oxygen and sulfur isotope values of CAS samples (e.g.

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Stratosphere Troposphere

Volcanic Emissions

Fig. 4. Schematic of the δ34S system and its interpretation within sulfate. Text in red indicates processes that will drive δ34S values more positive, text in blue indicates processes that will drive δ34S values more negative, and yellow indicates a process that can either drive values more positive or negative or does not impart any significant fractionation.

+

δ34S Sulfide Weathering Evaporite Weathering/ Deposition Microbial S Cycling Hydrothermal S-Reoxidation Abiotic S Reduction Sulfide burial

analytical procedure are 0.1‰ for δ34S measurements and 0.01‰ for ∆33S measurements.

SF6 gas was purified through multiple cold-traps under vacuum, followed by gas chromatography. The purified samples were analyzed in dual inlet mode on a Thermo MAT-253 in the McGill Stable Isotope Laboratory. Results were measured against international standard reference material IAEA-S1 (δ34S = –0.3‰ and Δ33S = –0.061). Estimated maximum errors (1σ) on measurements and the entire

4.4. GPP calculations To estimate Proterozoic GPP from our ∆17O measurements we Fig. 5. Schematic of controls on the ∆33S composition of sulfate. Text in red indicates processes that will drive Δ33S values more positive, text in blue indicates processes that will drive Δ33S values more negative, and yellow indicates a process that can either drive values more positive or negative or does not impart any significant fractionation.

Volcanic Emissions

+

Weathering Archean Terranes (S-MIF?)

∆33S

Sulfide Weathering Evaporite Weathering/ Deposition Microbial S Cycling Hydrothermal S-Reoxidation Abiotic S Reduction

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Fig. 6. Map of sulfate evaporites sampled. Locations of samples analyzed in this study are presented as different coloured stars separated into different time intervals and labeled with numbers corresponding to information in Table 1. Earliest Proterozoic (GOE) locations are in dark blue; mid-Proterozoic locations are in green; late Proterozoic locations are in red; Cryogenian locations are in light blue; Ediacaran locations in purple; and modern/Cenozoic samples in cream.

employed methods and assumptions outlined in Crockford et al. (2018), that are based on initial modeling work by Cao and Bao (2013). In this approach we combine ∆17O values in sulfate together with constraints on ancient atmospheric chemistry and other atmospheric parameters (e.g. stratosphere-troposphere exchange). This approach allows for calculations of the gross oxygen flux from the biosphere to the atmosphere which we approximate as GPP. In this approach the ∆17O value of tropospheric O2 is sensitive to pCO2, pO2, stratosphere-troposphere exchange, stratospheric photochemistry and GPP. We calculate the ∆17O values of Proterozoic tropospheric O2 by applying laboratory calibrations of sulfide oxidation that constrains the atmospheric oxygen component of product sulfate to between 8 and 15% in natural settings (Balci et al., 2007; Percak-Dennett et al., 2017). We apply the same constraints for other model parameters as Crockford et al. (2018) with the exceptions of pCO2 and pO2 where we use currently available estimates provided across the entire Proterozoic as summarized in Table 2 and the upper panel of Fig. 9. Since all subsequent reactions involving sulfate-bound-oxygen will move ∆17O values more positive, it is difficult to disentangle such processes from true stratigraphic variability. Therefore, we utilize all ∆17O values in GPP calculations to provide the full range of possible solutions (light green field, panel B Fig. 9), but also provide alternative possible GPP trajectories (grey and dark green fields, panel B Fig. 9). In Fig. 9 panel B the grey field represents possible solutions within the 25th and 75th percentiles inputting all suggested pO2 estimates and assuming all ∆17O values represent true stratigraphic variability. In the dark green field of Fig. 9 panel B we plot values assuming the most negative ∆17O value represents the most un-altered atmospheric signature and combine this with a lower O2 trajectory (dark blue region Fig. 9 panel A). This lower GPP trajectory (dark green field, panel B Fig. 9) also takes into account that the full range of ∆17O values observed in a single formation may be an unlikely range of pO2-pCO2-GPP variability over the time period captured by the stratigraphic interval between sulfate minerals. We present estimates of GPP as probabilities for each time interval based on a Monte Carlo approach where calculations were made using

model parameters randomly chosen from defined ranges and type of distribution (i.e. uniform versus Gaussian) and repeated 10,000 times. These calculations provide a conservative range of GPP where the upper bound of the uncertainty envelopes in the middle panel of each trajectory in Fig. 9 panel B, represents a high O2 – low CO2 atmosphere and ∆17O values closest to 0‰ (light green and grey trajectories), and lower bounds represent a low O2 – high CO2 atmosphere and the most negative ∆17O value for a given time interval. 5. The isotopic record of sedimentary sulfate In this section we discuss the isotopic record of sedimentary sulfate from the Archean through to the modern with an emphasis on the Proterozoic Eon, which we subdivide into five different intervals (Figs. 7 and 9) that are discussed in separate sections. Given that large age uncertainties are endemic to evaporite deposits, together with significant gaps in the sulfate evaporite record, we set these boundaries with an eye toward highlighting changes in the sulfate isotope record. We also consider other geochemical records sensitive to Earth's surface evolution such as changes in the δ13C of carbonates and organic matter that have been interpreted to indicate variations in the fraction of carbon buried as organic matter through time (forg; Des Marais et al., 1992; Hayes and Waldbauer, 2006; Krissansen-Totton et al., 2015; Fig. 7). These boundaries are broadly consistent with previously suggested changes in the redox state of the Earth's atmosphere and oceans (Cloud, 1972; Rouxel et al., 2005; Scott et al., 2008; Holland, 2006; Lyons et al., 2014; Havig et al., 2017). 5.1. The Archean Archean surface environments were unlike Proterozoic or Phanerozoic ones with profoundly different atmospheric and marine chemistry and redox state. In the low-oxygen Archean, UV radiation reached the lower atmosphere and drove sulfur photochemical reactions imparting large S-MIF signatures to product sulfur species that, for the most part, have not been preserved since this interval of Earth 208

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Fig. 7. Isotopic record of Proterozoic sulfate. New results from this study (red circles = evaporites; blue triangles = CAS; peach squares = barites) compiled with previously published δ34S, δ18O, ∆33S, and ∆17O data (grey circles = evaporites; grey squares = barites; grey triangles = CAS). Data is compiled from: Azmy et al., 2001; Bao et al., 2008, 2009, 2012; Blättler et al., 2018; Burke et al., 2018; Cameron, 1983; Chu et al., 2007; Claypool et al., 1980; Cortecci et al., 1981; Cowie and Johnston, 2016; Crockford et al., 2016, 2017, 2018; Das et al., 1990; Deb et al., 1991; Fike and Grotzinger, 2008; Fox and Videtich, 1997; Gellatly and Lyons, 2005; Gill et al., 2007; Goldberg et al., 2005; Grinenko et al., 1989; Guo et al., 2009; Halverson and Hurtgen, 2007; Holser and Kaplan, 1966; Hough et al., 2006; Hurtgen et al., 2002, 2004, 2005, 2009; Johnston et al., 2005b; Kah et al., 2004, 2016; Kampschulte and Strauss, 2004; Kampschulte et al., 1998; Kesler and Jones, 1980; Killingsworth et al., 2013, 2018; Krupenik et al., 2011; Li et al., 2015; Longinelli and Flora, 2007; Luo et al., 2015; Markovic et al., 2016; Masterson et al., 2016; Mazumdar and Strauss, 2006; Master et al., 1993; Misi and Veizer, 1998; Ortí et al., 2010; Palmer et al., 2004; Paytan et al., 1998; Peng et al., 2011; Peryt et al., 2010; Pierre and Rouchy, 1986; Planavsky et al., 2012; Present et al., 2015; Rennie et al., 2018; Reuschel et al., 2012; Rick, 1990; Ries et al., 2009; Sakai, 1972; Schobben et al., 2017; Schröder et al., 2008; Shi et al., 2018; Shields et al., 2004; Sim et al., 2015; Strauss, 1993, 1999; Strauss et al., 2001; Surakotra et al., 2018; Thompson and Kah, 2012; Tostevin et al., 2014, 2017; Turchyn and Schrag, 2004, 2006; Turchyn et al., 2009; Ueda et al., 1987; Ueda et al., 1991; Utrilla et al., 1992; Velikoslavinsky et al., 2003; Worden et al., 1997; Wotte et al., 2012; Wu et al., 2010, 2014, 2015; Yang et al., 2018; Yao et al., 2018 and new data from this study. Please refer to individual studies for associated errors on analysis and refer to the Methods section of this paper for errors on newly generated data. In the left portion of the figure we plot summary statistics where the white bars represent median values, the darker coloured boxes represent solutions within 25th and 75th percentiles of results and the light coloured boxes represent the total range of data.

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Table 1 Formations, locations and ages of samples used in this study. Evap* indicates hotspring and modern weathering product sulfate samples. Map #

Region

Modern/Cenozoic 1 N. Canada 2 SW. USA 3 SW. USA 4 SW. USA 5 S. Australia 6 Spain 7 N. Canada

Area

Unit

Age (Ga)

SO4 type

References

Axel Heiberg I. California California Nevada Flinders R. Sorbas B. Devon I.

N/A N/A N/A N/A N/A N/A N/A

0 0 0 0 0 0.006 0.014

Evap* Evap* Evap* Evap* Evap Evap Evap*

Oskoba

Evap Evap CAS

Khomentovsky, 2008; Mel'nikov, 2005, 2009 Houghton, 1980; Strauss, 1993

Ediacaran 541–635 Ma 8 Siberia 9 Iran 10 N. Australia

Kimberley

Egan

0.56 0.545 0.58

Cryogenian 635–717 Ma 11 N. Australia

Kimberley

Landrigan

0.635

CAS

Condon et al., 2005

late-Proterozoic 717–1100 Ma 12 NW. Canada Mackenzie Mts 13 N. Canada Victoria I. 14 N. Canada Brock Inlier 15 S. Australia Flinders R. 16 C. Australia Amadeus B. 17 NW. Canada Mackenzie Mts 18 W. Australia Officer B. 19 Zambia 20 D. R. Congo 21 N. Canada Victoria I. 22 N. Canada Baffin I.

Redstone R. Kilian Fm. Minto Inlet Fm. Skillogollee Loves Creek Fm. Ten Stone Browne Roan Mbuji/Mayi Minto Inlent Angmaat

0.75 0.8–0.72 0.795 0.8 0.8 0.81–0.89 0.83 0.883 (0.893–0.873) 0.883 0.85–0.89 1.05

Evap Evap Evap CAS Evap Evap Evap Evap Evap Evap Evap

Jefferson and Parrish, 1989 Rayner and Rainbird, 2013 Rayner and Rainbird, 2013 Turner and Bekker, 2016 Hill and Walter, 2000; Preiss, 2000 Armstrong et al., 2005 Cahen, 1982; Delpomdor et al., 2013 Van Acken et al., 2013; Rayner and Rainbird, 2013 Gibson et al., 2017

mid-Proterozoic 1100–2000 Ma 23 N. Australia McArthur B. 24 E. India Cuddapah 25 Canada Sibley

Myrtle Shale Vempalle Rossport

1.7 1.89 1.4

Evap Evap Evap

Muir, 1987; Walker, 1977; Kunzmann et al., 2019 Collins et al., 2015; French et al., 2008 Crockford et al., 2018

GOE 2000–2450 Ma 26 N.W. Russia 27 Siberia, Russia

Karelia Aldan Shield

Tulomozero Fedorovka

2.09 (2.16–2.02) or 1.98 2.1

Evap Evap

28 29 30 31 32 33

Griqualand West Basin Magondi Basin Yerrida Deweras Ontario Transvaal basin

Lucknow Lomagundi Juderina Norah Gordon Lake Duitschland

2.16 (1.86–2.46) 2.15 (2.1–2.2) 2.173 (2.237–2.109) 2.262 2.308 (2.3–2.316) 2.4

Evap Evap Evap Evap Evap Evap

Ovchinnikova et al., 2007; Kuznetsov et al., 2010; Martin et al., 2015 Vinogradov et al., 1976; Zolotarev et al., 1989; Velikoslavinsky et al., 2003 Schröder et al., 2008 Schidlowski and Todt, 1998; Master et al., 2010 Woodhead and Hergt, 1997; Sheppard et al., 2016 Manyeruke et al., 2004; Master et al., 2010 Chandler, 1988; Rasmussen et al., 2013 Bekker et al., 2001; Gumsley et al., 2017

S. Africa Zimbabwe W. Australia Zimbabwe S. Canada S. Africa

history (Farquhar et al., 2000; Farquhar and Wing, 2003; Endo et al., 2016; Ono, 2017; Eickmann et al., 2018). Although the record of Archean sulfate deposits is sparse, existing ∆17O data for ca. 3.5 Ga barites from Western Australia and 3 2 Ga barites from South Africa show no OMIF (Farquhar et al., 2000; Bao et al., 2007; Crockford et al., 2018). A lack of O-MIF suggests low oxygen levels below the suggested 10−3 PAL threshold required to build up a significant ozone layer that facilitates O-MIF production through photochemical reactions (Kasting and Donahue, 1980; Segura et al., 2003). Moreover, the existence of large SMIF at this time provides additional constraints on the Archean atmosphere. Although the chemistry involved in the production of S-MIF is complicated and debated (Babikov, 2017; Harman et al., 2018), Pavlov and Kasting (2002) calculated that the S-MIF signature would likely only exist in an atmosphere with extremely low ozone concentrations, corresponding to pO2 levels between 10−5 and 10−13 PAL. A lack of an ozone layer would have permitted short wavelength radiation to penetrate the atmosphere, potentially limiting environments where early life could survive (Mloszewska et al., 2018). Although it remains debated when oxygenic photosynthesis evolved (Buick, 1992; Des Marais, 2000; Farquhar et al., 2011; Planavsky et al., 2014a, 2014b; Shih et al., 2017; Wang et al., 2018), isotopic evidence for low O2 supports either the absence or extreme restriction of oxygenic photosynthesis across the Archean surface environment. If present at all, such low levels of oxygenic photosynthesis would have

resulted in a much smaller biosphere and hence much lower GPP on the Archean Earth compared to Proterozoic or Phanerozoic Eons (Kharecha et al., 2005; Canfield et al., 2006; Ward et al., 2016, 2018). While Archean atmospheric chemistry resulted in large mass-independent isotopic signatures of sulfur, but not oxygen (Farquhar et al., 2000; Bao et al., 2007; Crockford et al., 2018), surface mass-dependent processes isotopically fingerprinted sulfur pools, and these signals have been preserved in the sedimentary record (Halevy et al., 2010; Halevy, 2013). Under an assumption that the pyrite record reflects primary microbial byproducts, large δ34S fractionations between barite and associated pyrites with similarly negative ∆33S values in 3.5 Ga barites have been interpreted to indicate that dissimilatory sulfate reduction was an important metabolism driving the sulfur cycle at this time (Shen et al., 2001; Ueno et al., 2008; Shen et al., 2009). While other metabolisms that characterize the modern Earth may have originated in Archean environments (Ueno et al., 2006; Hug et al., 2016), the degree to which they operated would likely have been far different in the absence of significant carbon fixation from oxygenic photosynthesis (Des Marais, 2000). These isotopic records are also consistent with much lower sulfate inventories (e.g., ≈80 μM Jamieson et al., 2013; < 2.5 μM Crowe et al., 2014; Horner et al., 2017) than across the Proterozoic and Phanerozoic Eons. Low sulfate levels were a likely consequence of extremely low pO2 and limited oxidation of reduced sulfur (Krouse, 1980; Kaplan, 1983). Although some similar major isotope signatures are 210

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Fig. 8. Cross plots of isotopic measurements (A ∆17O-δ18O; B ∆17O-δ34S; C ∆17O-∆33S; D δ18O-δ34S; E ∆33S-∆36S; F ∆33S-δ34S) from this and previous studies. The field outlined in grey represents the predicted range of possible values from a steady state global sulfur cycle model (Johnston et al., 2005b) augmented to include an expanded range of possible 34ε (or 34α) and 33λ (or 33θ) from Sim et al. (2015). Table 2 Input pCO2 and pO2 values for GPP calculations for Fig. 9. CO2 and O2 values are given as PAL (CO2 1 PAL = 280 p.p.m.; O2 1 PAL = 209,500 p.p.m.). CO2 estimates are taken from the literature where we averaged upper and lower estimates provided (Mills et al., 2014; Fiorella and Sheldon, 2017; Wolf and Toon, 2014; Sheldon, 2006; Rye et al., 1995; Retallack and Mindszenty, 1994; Zbinden et al., 1988; Sheldon, 2013; Holland et al., 1989; Wiggering and Beukes, 1990; Kah and Riding, 2007; von Paris et al., 2008; Blättler et al., 2017; Kanzaki and Murakami, 2015). O2 estimates are also taken from the literature (Planavsky et al., 2014b; Bachan and Kump, 2015; Canfield, 2005; Blamey et al., 2016; Cole et al., 2016). Age (Ma) Trajectory All possible solutions

High O2, all ∆17O values

Low O2, min. ∆17O values

2325 2150 1890 1700 1400 1050 880 810 750 560 2325 2150 1890 1700 1400 1050 880 810 750 560 2325 2150 1890 1700 1400 1050 880 810 750 560

pO2 (PAL)

pCO2 (PAL)

GPP (Modern)

O2 max

O2 min

CO2 max

CO2 min

2.5th percentile

2 2 0.1 0.1 0.1 0.1 0.1 0.5 0.5 0.5 1 1 0.1 0.1 0.1 0.1 0.1 0.5 0.5 0.5 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.1 0.1 0.1

0.1 0.1 0.001 0.001 0.001 0.001 0.001 0.01 0.01 0.01 0.1 0.1 0.001 0.001 0.001 0.001 0.001 0.01 0.01 0.01 0.001 0.001 0.001 0.001 0.001 0.001 0.001 0.01 0.01 0.01

80 100 50 40 30 17 14 12 12 20 80 100 50 40 30 17 14 12 12 20 80 100 50 40 30 17 14 12 12 20

10 8.5 6.5 5 2 1 2 2 1 1 10 8.5 6.5 5 2 1 2 2 1 1 10 8.5 6.5 5 2 1 2 2 1 1

0.079 0.078 0.034 0.018 0.016 0.055 0.055 1.00 0.27 0.33

211

25th percentile

6.21 6.47 0.13 0.067 0.064 0.19 0.21 0.37 0.34 0.42 0.074 0.081 0.040 0.029 0.025 0.070 0.059 0.059 0.50 0.56

Median

10.3 10.7 0.42 0.21 0.18 0.46 0.50 0.79 0.74 1.03 0.13 0.15 0.075 0.055 0.046 0.12 0.11 0.11 0.77 0.88

75th percentile

16.8 17.2 1.11 0.54 0.44 0.99 1.04 1.46 1.40 2.13 0.25 0.27 0.14 0.10 0.084 0.21 0.18 0.18 1.14 1.39

97.5th percentile 27.3 35.4 3.68 1.51 1.36 2.77 3.03 5.45 3.58 6.92

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Fig. 9. A calculated record of gross primary production (GPP) across the Proterozoic eon. Calculations are based upon Proterozoic pCO2 (orange curve; panel A), and pO2 estimates (purple curve, blue = low O2 curve; panel A) together with ∆17O values of sulfate calculated through the model of Cao and Bao (2013) applied in the same way as Crockford et al. (2018). Calculated GPP values are plotted in panel B where the light green envelope encompasses the total possible range in Proterozoic estimates together with previous Archean estimates (Kharecha et al., 2005; Canfield et al., 2006; Ward et al., 2016, 2018) and Phanerozoic estimates from Wing (2013). The dark green and grey curves represent 25th and 75th percentiles of solutions with the grey field representing results that allow all ∆17O values and the full pO2 range from the literature (purple field, panel A). The dark green field only considers the most negative ∆17O value of a given time period and lower pO2 estimates (blue curve, panel A). The lower panel shows calculated forg values from Krissansen-Totton et al. (2015) to compare our model results for GPP with independent measures of the ancient carbon cycle.

shared by modern, post-GOE and Archean sulfates (Shen et al., 2001), the relatively large contribution of an atmospheric sulfur flux at this time prevents direct calibration and comparison of pre-GOE with postGOE sulfur isotopic signatures. These factors make the Archean sulfur and oxygen cycles geochemically distinct from any later period in Earth history and thus Archean sulfate isotope records can serve as a baseline in interpreting later records.

8). The present study provides the first confirmation of predicted mirrored mass-independent ∆33S - ∆17O records across the GOE (Bao, 2015; Fig. 7). Moreover ∆17O results provide a new constraint on minimum pO2 at this time, as the inception of O-MIF, infers that pO2 increased above sufficient levels to establish a stable ozone layer (pO2 > 10−3 PAL; Kasting and Donahue, 1980; Segura et al., 2003; Goldblatt et al., 2006). Since present day O2 with a ∆17O value of ≈−0.5‰ (Pack et al., 2017) cannot impart more than ≈30% of its signature into product sulfate (Balci et al., 2007; Kohl and Bao, 2011; Cao and Bao, 2013), GOE values as low as −0.36‰ must reflect different environmental conditions than the modern Earth. The implication is that observed ∆17O values are unlikely to be completely reset by contamination from modern sulfate or isotope exchange with water earlier in its history and can therefore provide insights into the biosphere and atmosphere at this time. Through new and previously published ∆17O results (Blättler et al., 2018), estimates of GPP across the GOE can be calculated provided that pO2 and pCO2 can be constrained (Cao and Bao, 2013; Crockford et al., 2018). When considering suggested estimates of pCO2 and pO2, it is important to note that they do not capture the likely dynamic nature of atmospheric chemistry across the GOE. Extensive glaciations (Young, 1991), possibly global in extent (Evans et al., 1997; Kirschvink et al., 2000; Kopp et al., 2005; Bekker, 2014a), engulfed the Earth and must

5.2. Earliest Proterozoic (GOE) (2.5–2.0 Ga) Results from this work begin with the first occurrences of sulfate evaporites at ca. 2.4 Ga that were deposited near the initiation of the GOE (Roscoe, 1973; Luo et al., 2015). The first-order observation from our sulfate isotope data-set is mirrored trends in O-MIF and S-MIF (i.e., S-MIF but no O-MIF before the GOE, O-MIF but no S-MIF during and after the GOE) that are consistent with the first sustained increase in atmospheric oxygen in the earliest Proterozoic (defined here as 2.5–2.0 Ga therefore comprising only the first-half of the Paleoproterozoic) (Macgregor, 1927; Farquhar et al., 2000; Bekker et al., 2004). Beginning with ∆17O values, an increase in the magnitude of O-MIF away from Archean values (at ca. 3 2 Ga ∆17O > −0.16‰ Bao et al., 2007; 3.5 Ga ∆17O > −0.14‰ Crockford et al., 2018) to early Proterozoic/GOE ∆17O values as low as −0.36‰ is observed (Figs. 7 and 212

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have induced large perturbations to the Earth system (cf. Bekker et al., 2005; Bekker and Kaufman, 2007; Zahnle, 2006; Konhauser et al., 2009). Here we explore ∆17O results through the long-lasting, nonglacial intervals across the GOE. The early Proterozoic Earth received less solar radiation from a dimmer Sun (Gough, 1981) implying that higher than modern levels of CO2 and possibly some other greenhouse gases (cf. Roberson et al., 2011; Stanton et al., 2018), were required to maintain ice-free surface conditions (Walker et al., 1981). Existing estimates of pCO2 that fall between ≈10 and 100 PAL across the GOE support this inference (where 1 PAL = 280 p.p.m.; von Paris et al., 2008; Kanzaki and Murakami, 2015; Wolf and Toon, 2014; Blättler et al., 2017; Rosing et al., 2010; Sheldon, 2006; Isson and Planavsky, 2018). Estimates of pO2 span a wider range and typify the uncertainty in the magnitude of this event, despite improved age constraints on its duration (e.g., Gumsley et al., 2017). The GOE encapsulates the Lomagundi-Jatuli positive carbon isotope excursion in sedimentary carbonates that globally records long lived δ13C values > 5‰ (Karhu and Holland, 1996; Bekker, 2014b). Notwithstanding arguments against the direct interpretation of carbon isotope values in carbonates as a direct measure of the relative burial rate of organic carbon (Immenhauser et al., 2003; Swart, 2008; Higgins et al., 2018; Ahm et al., 2018), the Lomagundi-Jatuli carbon isotope excursion has been interpreted to be a result of extraordinarily high organic carbon burial rates that may have resulted in a large flux of oxygen to the atmosphere (Des Marais et al., 1992; Karhu and Holland, 1996). These inferences have led to the idea that the GOE represents a period of an ‘oxygen overshoot’ where pO2 not only rose above Archean levels but was characterized by higher pO2 than the subsequent mid-Proterozoic (Bekker and Holland, 2012; Bachan and Kump, 2015; Planavsky et al., 2018). Although this interpretation is supported by shifts in various geochemical proxies across the GOE (e.g., Scott et al., 2008; Partin et al., 2013), it remains difficult to constrain pO2 across this interval beyond lower limit O-MIF constraints (minimum pO2 = 0.1% modern (Segura et al., 2003) and upper limit estimates based on model calculations (e.g. ≈200% modern; Bachan and Kump, 2015). Here we explore this interval with a wide range of pO2 estimates between 0.1% to 200% modern and input the full suite of measured ∆17O values for GPP calculations (light green field, Fig. 9 panel B). We then explore the GPP consequences of such high pO2 estimates with alternative pO2 paths along with only utilizing the most negative ∆17O measurement from a given interval (dark green curve, Fig. 9 panel B). Given the ranges presented above, we calculate GPP across the GOE to be between over 1000% to < 10% modern with a median value of ≈300% modern (Fig. 9). Importantly, the ∆17O record within sedimentary sulfates can only capture a fraction of the atmospheric signal since any deposited sulfate bears some history of alteration away from initial oxidation products and thus provide maximum GPP estimates. Nevertheless, our results suggest the GOE may have potentially been characterized by the highest sustained levels of GPP through all of Earth history. Such high GPP estimates are a direct consequence of lower estimates of pCO2 and, more importantly, of extremely high pO2 estimates (Fig. 9). While none of the isotopic systems within sulfate can directly contradict such high pO2 estimates, results of our GPP calculations highlight the biological consequences of such a prediction. That is, if such high GPP levels sustained over > 100 Myrs cannot be reconciled within biogeochemical models that explore ancient nutrient cycling, ∆17O values presented here may suggest that the upper end of pO2 estimates should be revisited (Bachan and Kump, 2015) and that GPP levels may be better represented by the dark green (lower GPP) trajectory plotted in Fig. 9. While the GOE may have been characterized by high levels of GPP, controversy has surrounded the early Proterozoic interval that bridges the gap between the GOE and the Archean. Beginning at ca. 2.4 Ga with Earth's earliest evaporites, ∆33S values fall to within the typical massdependent range that continues through to measurements of sulfur

isotopes in modern environments (Farquhar and Wing, 2003). Whether a result of an increase in organic carbon burial due to the advent of oxygenic photosynthesis (Soo et al., 2017), a rise to dominance of this metabolism millions of years after its emergence (Castresana and Saraste, 1995; Pereira et al., 2001; Crowe et al., 2013; Planavsky et al., 2014a; Wang et al., 2018), tectonomagmatic evolution linked to the assembly and breakup of supercontinents/supercratons (Gumsley et al., 2017), or a sharp decrease in reductant fluxes to surface environments (Ebelmen, 1845; Berner and Maasch, 1996; Canil, 1997; Kump et al., 2001), the disappearance of S-MIF signatures from the sedimentary record indicates an increase in atmospheric oxygen levels above ~10−5 PAL (Farquhar et al., 2000; Pavlov and Kasting, 2002; Guo et al., 2009). Evidence for early oxygenation through coupled S-MIF and O-MIF trends is consistent with observable macroscale features of the sedimentary record. Such features include: the disappearance of detrital pyrites and uraninites (Rasmussen and Buick, 1999), the formation of large manganese deposits (Laznicka, 1992; Kirschvink et al., 2000; Kunzmann et al., 2014), the appearance of redbeds, and perhaps most notably, the first appearance of sulfate evaporites at ca. 2.4 Ga that indicate marine sulfate concentrations above ≈2.5 mM (Wood, 1973; Bekker et al., 2006; Schröder et al., 2008). Some studies have raised concerns about drawing direct links to the texture of the S-MIF record and the pace of oxygenation of the surface environment observed through these features outlined above (Philippot et al., 2018; Reinhard et al., 2013a). For example, S-MIF has been reported in Proterozoic terranes recycled from older Archean rocks (Selvaraja et al., 2017). Others have argued that initial oxidative terrestrial pulses bearing SMIF signatures may extend the S-MIF sedimentary record well after these signals cease to be produced in the atmosphere (Reinhard et al., 2013a). This S-MIF-recycling hypothesis has found recent support through S-MIF occurences within sediments from a glacial-bearing succession in Australia (Turee Creek Group; Philippot et al., 2018). However, the lack of S-MIF in the new data presented here from Earth's earliest sulfate evaporite deposits suggests that such recycling was not able to exert a measurable influence on marine sulfate as of ca. 2.4 Ga (Fig. 7). This finding is also consistent with recent results showing no SMIF preserved in waters derived from weathering Archean watersheds (Torres et al., 2018). That is, the marine sulfate ∆33S record as preserved within sulfate evaporites lends credence to suggestions that the accumulation of oxygen in the atmosphere and subsequent increase in the intensity of sulfur cycling was relatively rapid once atmospheric production of S-MIF ceased (Bekker and Kaufman, 2007; Gumsley et al., 2017; Luo et al., 2016). Moreover, analyses of over 80 samples from seven different formations on three different continents, all yielding ∆33S values without mass-independent signatures, are inconsistent with globally and temporally significant recycling of Archean S-MIF (Reinhard et al., 2013a; Philippot et al., 2018). Isotopic patterns observed through the minor isotopes and their concomitant interpretation can be tested against isotopic trends in the major isotopes of sulfate. Major sulfur isotope (δ34S) data range between 4 and 42‰, but most data cluster near an average of 16‰ that is slightly lower than that of modern seawater (21.2‰; Tostevin et al., 2014; Fig. 7). The range of δ34S values compared to the mid-Proterozoic is significantly smaller (Fig. 7). Less variation in δ34S across the GOE compared to that in the mid-Proterozoic is consistent with a comparatively larger marine sulfate reservoir at this time (e.g., Bekker et al., 2006; Schröder et al., 2008; Blättler et al., 2018), an inevitable consequence of higher atmospheric oxygen concentrations. Most values cluster around a weak trend of decreasing δ34S values across the GOE (Fig. 7) that possibly reflects increased pyrite burial and a concomitant contraction of the marine sulfate reservoir (Planavsky et al., 2012; Scott et al., 2014). Although sulfur isotope trends can be explained by local processes (e.g., Paiste et al., 2018), the expression of the δ18O record is also likely to be a response to a contraction of the marine sulfate reservoir at the end of the GOE, however identifying direct drivers of observed trends is challenging. 213

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Highly positive δ18O values are observed across the GOE with a maximum value of 36.1‰ and an average value of 13.1‰ (approximately 4‰ heavier than modern seawater sulfate; Brand et al., 2009; Johnston et al., 2014) Fig. 7; Table 1). Notwithstanding extreme outliers, a plausible global mechanism to drive such positive average δ18O and δ34S values could be vigorous dissimilatory sulfate reduction coupled to high iron availability to sequester product sulfide as pyrite in sediments, which would leave residual sulfate isotopically heavy (Wortmann et al., 2007). This mechanism requires a large supply of organic substrate to fuel sulfur cycling, which may indirectly reflect the changes in forg and GPP required to drive the Lomagundi-Jatuli carbon isotope excursion (Bekker and Holland, 2012; Bekker, 2014b; Krissansen-Totton et al., 2015; Fig. 9). In sum, the GOE saw an irreversible increase in pO2 levels and a growth in the marine sulfate reservoir that is borne out in the disappearance and appearance of ∆33S and ∆17O anomalies, respectively, and the deposition of Earth's first sulfate evaporite deposits that preserve these signatures. While broadly consistent with suggested trajectories in the redox state of the surface Earth (Lyons et al., 2014), additional details over this interval are required to understand the full complexity of sedimentary sulfate isotope records. Although extreme levels of GPP are permissible with ∆17O results when interpreted together with existing atmospheric estimates, it is unclear whether such high levels of primary productivity are feasible given the high nutrient fluxes or extreme levels of biomass recycling that would be required to sustain such a biological state.

in the Earth System. Irrespective of the GPP trajectory in Fig. 9, if the range of ∆17O values in sulfate deposits has captured true minimum values produced during terrestrial pyrite oxidation, trends between deposits may reflect a secular trend in the size of the biosphere and composition of the atmosphere (Fig. 9) that possibly record changing nutrient dynamics over this nutrient limited mid-Proterozoic interval (Canfield, 1998; Anbar and Knoll, 2002; Robbins et al., 2016; Koehler et al., 2017; Ozaki et al., 2019). Changing nutrient inventories over the Proterozoic has been suggested as a mechanism to both temper and invigorate the biosphere (Laakso and Schrag, 2014; Reinhard et al., 2013b; Derry, 2015; Sánchez-Baracaldo et al., 2014; Cox et al., 2016a, 2016b; Reinhard et al., 2016; Kuznetsov et al., 2017; Koehler et al., 2017). With studies suggesting decreasing pCO2 levels over the Proterozoic, one would predict a concomitant shift in ∆17O values toward ≈0‰ if all other factors were held constant (Cao and Bao, 2013; von Paris et al., 2008; Sheldon, 2013). However, the opposite trend is observed between ca. 1.9 and 1.4 Ga (i.e. progressively more negative ∆17O values). The only viable culprit is low GPP that outcompeted decreasing atmospheric pCO2 levels in levering ∆17O values. Importantly, GPP estimates from this work (100 > GPP > 10−2 modern; Fig. 9) appear to be consistent with a recent attempt to explore the net primary productivity of the Proterozoic biosphere (Laakso and Schrag, 2019) and this previous work is most consistent with the lower GPP trajectory (dark green, panel B) in Fig. 9. Moreover, our results suggest that GPP across the GOE-mid-Proterozoic transition may have been one of the largest sustained changes to the biosphere across all of Earth history. This observation even holds true under a scenario where pO2 is held constant across the GOE and mid-Proterozoic. It stands to reason that a weak mid-Proterozoic biosphere would ultimately result in lower atmospheric oxygen concentrations than the modern Earth, assuming that reductant fluxes to the surface environment as well as forg values fall within previously estimated ranges (Sleep and Zahnle, 2001; Krissansen-Totton et al., 2015). Low oxygen efflux from the biosphere over this interval is consistent with previous studies that suggest low mid-Proterozoic atmospheric oxygen levels (Planavsky et al., 2014b; Cole et al., 2016; Zhang et al., 2016; Liu et al., 2016; Daines et al., 2017; Holland et al., 1989; Crockford et al., 2018; Planavsky et al., 2018). Low atmospheric and marine oxygen concentrations may have also led to a decrease in the size of the marine sulfate reservoir via increased iron availability and enhanced burial of pyrite in anoxic and sulfidic settings (Meyer and Kump, 2008; Lyons et al., 2009; Planavsky et al., 2011; Gomes and Johnston, 2017), that possibly outpaced the burial of organic matter (Ozaki et al., 2019). Low seawater sulfate levels would potentially create a more dynamic system with respect to δ34S values, as the residence time of sulfate in the ocean would likely be considerably shorter compared to the modern environment and more prone to transient changes (Richter and Turekian, 1993; Johnston et al., 2006). Data presented and compiled here potentially support such a model, with a wider range in midProterozoic δ34S values compared to the GOE and later in the Proterozoic (Fig. 7). The extremes of this scatter, however, are almost exclusively preserved within CAS (Fig. 7), which is likely more susceptible to post-depositional modification during diagenesis (Fichtner et al., 2017; Ahm et al., 2018). Moreover, additional stratigraphically resolved records with tight age constraints are needed to conclusively support the existence of a small sulfate reservoir based solely on δ34S data, and thus requiring other lines of evidence. Although there is far less ∆33S data compared with δ34S data for mid-Proterozoic sulfates, a small sulfate reservoir is also consistent with observed low ∆33S values. Such values are suggestive of a weak sulfur cycle with respect to dissimilatory sulfate reducers, which would have been the dominant sulfur metabolism at this time (Kunzmann et al., 2017; Fig. 7). Moreover, lower mid-Proterozoic water-column redox and weaker microbial sulfur cycling may also explain isotopically light δ18O values in sulfate with limited opportunities for exchange with

5.3. The mid-Proterozoic 2.0–1.1 Ga The mid-Proterozoic (defined here as 2.0–1.1 Ga therefore comprising the latter half of the Paleoproterozoic and majority of the Mesoproterozoic) stands out from other times in Earth history with respect to sulfate isotope records in that it is characterized by consistently anomalous negative ∆17O values and low ∆33S and δ18O values. These patterns coupled with the greatest Proterozoic variation in δ34S values distinguish the mid-Proterozoic record from earlier or later intervals. These signatures however, are from evaporites only, preserved within a limited number of basins, none of which were deposited as thickly bedded sediments like many younger and older deposits (Blättler et al., 2018; Turner and Bekker, 2016). The mid-Proterozoic ∆17O record of sulfate bears some of the most negative ∆17O values outside of the Cryogenian. Three formations from India (1.89 Ga), Australia (1.7 Ga), and previously published results from Canada (1.4 Ga) all display highly negative minimum values of −0.59‰, −0.78‰ and −0.88‰, respectively (Fig. 7). Such negative ∆17O values in conjunction with existing pCO2 and pO2 estimates have been cited as evidence for a less productive mid-Proterozoic biosphere, which would have throttled oxygen export from surface waters to the troposphere (Crockford et al., 2018). Expanding reported anomalous ∆17O values from one to now three formations suggest that this signature of reduced GPP is likely robust and characteristic of the broader (2.0 Ga–1.1 Ga) mid-Proterozoic interval (Fig. 9; Crockford et al., 2018; Ozaki et al., 2019). Inputting existing pCO2 and pO2 estimates for the mid-Proterozoic from Table 2 and following methods for GPP calculations outlined above, suggests that regardless of trajectory plotted in Fig. 9, this interval of Earth history was characterized by lower sustained GPP than the modern or the preceding GOE interval. Although GPP in excess of modern values are probabilistically permissible outliers under a combination of maximum pO2 and minimum pCO2 conditions (Fig. 9), such values also derive from the least negative ∆17O measurements that have potentially been subject to microbially-mediated oxygen isotope exchange with water, leading to resetting away from initial oxidation products. Moreover, in some cases, it is difficult to envisage such extreme changes in ∆17O values over the time interval likely captured by sulfates in a given stratigraphic sequence being reflective of corresponding extreme changes in GPP-pCO2-pO2 conditions 214

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other oxygen-bearing species, which potentially produced more positive values observed over the GOE (Fig. 7). Alternatively, low δ18O values may also be due to a higher proportion of sulfate in mid-Proterozoic deposits being a result of direct oxidation (Turchyn et al., 2009). This interpretation also supports a low ocean redox state and low seawater sulfate levels as preserving such signals is more likely in such environments. Low mid-Proterozoic sulfate levels are also supported by observations from the sedimentary record along with the sulfur isotope record presented above (Fig. 7). The mid-Proterozoic has relatively few sulfate occurrences compared to adjacent intervals (e.g. Kah et al., 2004; Scott et al., 2014) and the few documented occurrences are limited in thickness. While preservation bias and secular variations in basin architecture or location are plausible, the observation of so few preserved sulfate evaporite deposits is most consistent with lower marine sulfate concentrations over this interval. Maximum ∆17O values also potentially provide insights into the mid-Proterozoic marine and terrestrial sulfate reservoirs. Sulfate within deposits from the Vempalle (Cuddapah basin), Myrtle Shale (McArthur basin), and Rossport formations (Sibley basin; Table 1) show low maximum values of −0.21‰, −0.54‰ and −0.35‰ that possibly indicate a much larger influence of continental runoff on the isotopic composition of the marine and terrestrial sulfate reservoirs (Figs. 2 and 7). Assuming that the ∆17O composition of seawater has not changed through Earth history, the above suggestion is consistent with sulfur isotope data that imply a much smaller mid-Proterozoic marine sulfate reservoir possibly with a much shorter residence time. Under these conditions it is likely that only local signals would be preserved in the sedimentary record (Richter and Turekian, 1993; Shen et al., 2002). In sum, this low-oxygen-low-sulfate condition outlined above was most likely a function of reduced mid-Proterozoic GPP, which is borne out by highly negative ∆17O values over this interval. These conditions likely dictated the preponderance and nature of sulfate evaporite deposits with no documented extensive thickly bedded occurrences; instead only heterolithic supratidal deposits with thin beds and nodules are present in the rock record.

et al., 2017). Therefore, although GPP estimates from our calculations appear relatively high compared to the modern or Phanerozoic (Wing, 2013) in a world with a small to absent terrestrial biosphere (Zhao et al., 2017), GPP estimates from ∆17O values can only provide maximum estimates and not minimum estimates, meaning lower GPP estimates may be more reflective of the late Proterozoic Earth (dark green trajectory Fig. 9). Furthermore, GPP results highlight the biological consequences of such high pO2 over this interval (light green and grey trajectories Fig. 9), lending an additional line of concern (cf. Yeung, 2017) regarding estimates that put forward such values (Blamey et al., 2016). Our ∆17O based GPP estimates are consistent with a growth in the productivity of the biosphere, as well as a rise in pO2 and resultant larger marine sulfate reservoir. An increase in the marine sulfate reservoir is supported by the reappearance of thickly bedded evaporites in the sedimentary record across the late-Proterozoic beginning at 1.05 Ga (Jackson and Cumming, 1981; Gibson et al., 2017, 2019), with some sequences up to a kilometer-thick (Aitken, 1981; Turner and Bekker, 2016). Sulfur isotopes lend an additional line of support for growth in the size of the late-Proterozoic marine sulfate reservoir through a decrease in the range of δ34S values compared to that in the mid-Proterozoic (Fig. 7), presumably because the sulfate reservoir was less prone to isotopic shifts from changes in input and output fluxes (Kah et al., 2004). ∆33S values also deviate from mid-Proterozoic values with most data plotting above zero (Fig. 7). This increase is part of a broader trend with more positive values across the GOE, shifting to more negative values over the mid-Proterozoic and returning to more positive values in the late Proterozoic (Fig. 7). The trend is consistent with a rise in GPP levels across the GOE, followed by a fall over the mid-Proterozoic and then a rise again over the late Proterozoic. Importantly, trends in GPP are observed even when pO2 is held constant between these intervals. These changes in primary production may have been manifested in increased organic matter supply to fuel microbial sulfate reduction, which would favor higher ∆33S values in the remaining sulfate pool. Inferences based upon ∆17O and sulfur isotope data can be tested with δ18O values. δ18O values within sulfate display a similar symmetric pattern to ∆17O and ∆33S results, with maximum values during the GOE and the late Proterozoic (average = 15.1‰) and lowest values over the mid-Proterozoic (Fig. 7). This trajectory of average δ18O values over the Proterozoic is a possible consequence of a variable degree of dissimilatory sulfate reduction and sulfide weathering, where the GOE and late Proterozoic experienced higher rates than the mid-Proterozoic that left the sulfate pool relatively isotopically enriched over these intervals (Wortmann et al., 2007). This scenario is fully consistent with increased atmospheric and marine oxygen levels resulting from increased primary production that would also provide both the organic substrate and increased sulfate levels to fuel microbial sulfate reduction (Fig. 9). An alternative mechanism to explain δ18O values is an increased opportunity for isotopic exchange in the atmosphere of sulfur species due to enhanced DMS production from algae that experienced a major diversification over this interval of Earth history (Kumar et al., 2002; Parfrey et al., 2011; Feulner et al., 2015; Gibson et al., 2017; Fig. 3) and likely contributed to the enhanced primary production of the biosphere. However, biomarker data suggest that algae may not have risen to ecological prominence until the late Cyrogenian (Brocks et al., 2017). Another possibility is that late Proterozoic deposits bear some portion of sulfate derived from locally weathered sulfate deposits that had an enriched δ18O value and subsequent cycling of this sulfate pushed values even higher. It is important to note, however, that ∆17O values as negative as −0.44‰ cannot be generated in modern evaporative settings, meaning that late Proterozoic sulfate must bear a portion of initial oxidation products, which in turn, supports their primary nature. Therefore, it seems that a change in microbial sulfur cycling prompted by increased GPP provides the most parsimonious explanation for the revealed trends from the GOE through to the later Proterozoic.

5.4. Late-Proterozoic (1.1–0.72 Ga) Sulfate isotope values over the late Proterozoic (defined here as the latest Mesoproterozoic to the immediate pre-Cryogenian), depart from observed trends over the mid-Proterozoic and are similar to values observed over the early Proterozoic; however, they were likely produced under different environmental conditions. ∆17O values for sulfate evaporite deposits beginning at ca. 1.05 Ga (Gibson et al., 2017, 2019) from the Angmaat Formation of Baffin Island and continuing to younger units from Northwest Canada, display minimum values down to −0.38‰ and −0.44‰ respectively. ∆33S values show a slight increase in the average value compared to mid-Proterozoic samples which is also observed in δ18O and δ34S values with averages of 15‰ and 23‰, respectively. Not only do ∆17O, ∆33S and δ18O values display similar patterns to GOE trends, these signatures are also often preserved within thickly-bedded evaporites that reappear in the sedimentary record after an apparent billion-year long hiatus. These patterns potentially reflect another growth in the size of the marine sulfate reservoir to a level approaching that attained across the early Proterozoic (Blättler et al., 2018). Estimates for pCO2 across the late Proterozoic are far lower than those for the GOE due to increasing solar luminosity (Gough, 1981; Table 2). Under lower pCO2 conditions and based on available pO2 estimates, GPP was potentially lower during the late Proterozoic than over the GOE (Fig. 9). As discussed earlier, ∆17O values of sulfates record a complex history from initial oxidation of sulfide on the continents to microbial processing and possible mixing with other sulfate sources before deposition. Additionally, the GPP calculations depend on estimates of atmospheric chemistry over this interval which vary widely (Blamey et al., 2016; Yeung, 2017; Planavsky et al., 2014b; Daines 215

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In sum, the record from 1.05 Ga up until the Cryogenian saw an increase in GPP along with likely increases in marine sulfate and atmospheric oxygen from the preceding mid-Proterozoic interval. Importantly, some combination of these and other factors likely created conditions prone to climatic instability and contributed to events that ultimately plunged the Earth into the Cryogenian (Hoffman et al., 1998; Hoffman et al., 2017). While the evaporite record only comes to within ca. 30 Myrs of the initiation of the Sturtian glaciation at ca. 717 Ma (Macdonald et al., 2010; MacLennan et al., 2018), ∆17O values from the ca. 750 Ma Redstone River and 790 Ma Kilian formations (Jefferson and Parrish, 1989) bear similar values to those produced in the modern environment. When combined with pCO2 and pO2 estimates, these compositions suggest that this time was characterized by relatively high GPP compared to the preceding mid-Proterozoic. The only other time in Earth history with sustained high GPP together with evidence for panglaciation is the earliest Proterozoic (Kirschvink et al., 2000; Kopp et al., 2005; Bekker, 2014a) and it remains an open question whether similar causal mechanisms initiated climatic instability for both time intervals. Other factors that may have contributed to the onset of Cryogenian glaciations include a decrease in the size of the dissolved inorganic carbon (DIC) reservoir and an increase in atmospheric oxygen levels (Kah et al., 2004; Planavsky et al., 2014b; Cole et al., 2016; Gilleaudeau et al., 2016). The former was at least partly a consequence of a progressive increase in solar luminosity requiring reduced atmospheric pCO2 to maintain clement surface conditions (Walker et al., 1981; Gough, 1981). Potentially consistent with these changes in the carbon cycle are high amplitude shifts in the carbon isotopic composition of shallow marine carbonates beginning with a positive carbon isotope shift in the Angmaat Formation in the Bylot Supergroup and continuing into the Bitter Springs negative carbon isotope excursion at ca. 810 to 790 Ma (Kah et al., 1999; Halverson et al., 2005, 2018; Macdonald et al., 2010; Swanson-Hysell et al., 2010, 2015) and the ca. 735 Ma Islay negative carbon isotope anomaly (Hoffman et al., 2012; MacLennan et al., 2018). Factors that may have contributed to a rise in atmospheric oxygen levels include a diversification of the biosphere with evidence for Earth's earliest sexually reproducing eukaryotes at ca. 1.05 Ga which may have increased export production (Butterfield, 2000; Tziperman et al., 2011; Gibson et al., 2017). These evolutionary steps may have increased the ratio of organic carbon to total carbon that is buried, consistent with previously inferred δ13C trends over this interval (Bartley and Kah, 2004; Krissansen-Totton et al., 2015). These changes to the biosphere may have resulted from changes in the global weathering budget that would have both increased phosphorus supply to the oceans, and decreased CO2 in the atmosphere (Cox et al., 2016b; Mckenzie et al., 2016; Horton, 2015). Finally, although the isotopic record within sulfate minerals cannot point to a definitive causal mechanism, a growth in marine sulfate may have indirectly contributed to initiating the Cryogenian. Volcanic eruptions through the Minto Inlet and Kilian Formations on Victoria Island Canada (Table 1), may have injected large quantities of sulfur aerosols into the stratosphere. This aerosol injection would have rapidly changed the radiation budget of the Earth, sufficiently destabilizing climate to trigger glaciation (Macdonald and Wordsworth, 2017), further underscoring that a rise in marine sulfate abundance may have foreshadowed Cryogenian glaciations.

Peng et al., 2011; Killingsworth et al., 2013; Benn et al., 2015; Crockford et al., 2016, 2017). The isotopic signatures are now reported from seven paleo-continents (Crockford et al., 2017) and have been interpreted as evidence for high syn-glacial pCO2 levels in the atmosphere that would have ultimately terminated the Marinoan glaciation (Bao et al., 2008; Cao and Bao, 2013). While ∆17O anomalies are now known to be non-unique to the Marinoan (Crockford et al., 2018), the appearance in the late Proterozoic of thick bedded evaporites that do not display large anomalous signatures, indicates that anomalous postMarinoan signatures are reflective of the conditions imposed by the Marinoan Snowball Earth and are not remobilized sulfate from midProterozoic deposits (Bao et al., 2008; Cao and Bao, 2013; Fig. 7). While similar atmospheric conditions experienced during the Marinoan glaciation are predicted for the older Sturtian glaciation, no sulfate deposits from immediate post-Sturtian glacial rocks have been found. This contrast between the post-Sturtian and post-Marinoan records suggests much different geochemical dynamics operating during the Sturtian glaciation, possibly a consequence of its much longer duration (ca. 59 Ma vs. > 4 Ma; Prave et al., 2016; Macdonald et al., 2010; Rooney et al., 2014, 2015; Hoffman et al., 2017; MacLennan et al., 2018). One manifestation of the differences between Sturtian and Marinoan glaciations is the restriction of iron formation deposition to the older event. These chemical sediments must have been deposited under ferruginous conditions, potentially indicating a larger hydrothermal or continental iron supply to the ocean and a significant decrease in the size of the marine sulfate reservoir due to reduced oxidative weathering of continental sulfides with respect to earlier Earth history (Cox et al., 2013, 2016a). It is predicted that sulfate inputs into a sub-glacial ocean would also be severely attenuated, although many outputs would also be affected, including microbial sulfate reduction due to limited organic substrates. The net effect of such conditions would be the drawdown of marine sulfate, which is consistent with ∆17O and δ34S data from post-Marinoan sequences (Hurtgen et al., 2006; Crockford et al., 2016). Therefore, the long duration of the Sturtian glaciation may have drawn down sulfate concentrations to levels too low to capture transient post-glacial signatures reflective of syn-glacial conditions. Alternatively, the consistent absence of transgressive systems tracts in post-Sturtian cap carbonate sequences raises the possibility of fundamentally different post-glacial dynamics of the sulfur cycle operating between the two Cryogenian glaciations (Crockford et al., 2017). While it is clear that the Cryogenian Earth was distinct from both earlier and later times, its impact on the evolution of the biosphere remains enigmatic, with life confined to specific niche environments (Hoffman, 2016) potentially serving as either an evolutionary “activated complex” or as a temporary bottle neck for the dramatic evolutionary changes observed in the biosphere over this interval (Javaux, 2007; Brocks et al., 2017). Although one can envisage dramatic changes to the sulfur and oxygen cycles over this interval of Earth history, the geochemical record remains too sparse to fully explore their relationship with the evolution of the biosphere. 5.6. The Ediacaran (0.635–0.542 Ga) The Ediacaran sulfur and oxygen cycles appear to have been different to any earlier time in Earth history. During this period, major evaporite deposits reappear in the sedimentary record after a hiatus across the Cryogenian that is consistent with a growth in the size of the marine sulfate reservoir from immediate post-Cryogenain levels (Hurtgen et al., 2002; Fike et al., 2006; Halverson and Hurtgen, 2007; Crockford et al., 2016) potentially after a short transient interval of pervasive euxinia and elevated pyrite burial (Lang et al., 2018). Minor isotopic values of sulfate display limited variability and values are comparable to those during the GOE (e.g., minimum ∆17O = −0.29‰ and average ∆33S = 0.02‰). Ediacaran major oxygen isotope data with an average δ18O value of 13.6‰ are similar to those of the late

5.5. The Cryogenian (0.72–0.635 Ga) Due to the unique climatic and geochemical conditions both during and in the immediate aftermath of two Snowball Earth glaciations, we separate the Cryogenian into its own stage even though sulfate data is relatively sparse over this interval of Earth history (Hoffman et al., 2017). Sulfate ∆17O data from the Cryogenian has only been generated from syn-glacial and immediately post-glacial Marinoan CAS and barite that both display highly negative values (Bao et al., 2008, 2009, 2012; 216

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Proterozoic and are slightly higher than that of modern marine sulfate. Major sulfur isotopes show some of the most positive δ34S values in Earth history (up to 34‰; Fig. 7), and these notably heavy values are also concomitant with the emergence of some coherent structure in this record with significantly less scatter than earlier times (Fig. 7). Although the Ediacaran isotopic record is similar to other times in Earth history, it occurs under the backdrop of dramatic changes to the Earth surface environment that include the emergence of the Ediacaran fauna and flora (Erwin et al., 2011) and changes to the sulfur cycle (Kunzmann et al., 2017). The growth in the marine sulfate reservoir (Crockford et al., 2016; Hurtgen et al., 2002, 2006; Halverson and Hurtgen, 2007) and inferred (and presumably coupled) increases in atmospheric oxygen levels (Sahoo et al., 2012), provide evidence that changes in the biosphere appeared in conjunction with changes in atmospheric and ocean chemistry. Although transient oxygenation events (Sahoo et al., 2016), disparate local records (Sahoo et al., 2012; Kunzmann et al., 2015; Miller et al., 2017), and ambiguities in the interpretation of proxy data (Sperling et al., 2015; Slotznick et al., 2018) may suggest a complex and perhaps protracted (Wallace et al., 2017) history of oxygenation over this interval, broad trends of an increased inventory of atmospheric oxygen appear to be borne out in the sulfur and oxygen isotope records of sulfate evaporites. ∆17O values within post-Marinoan barites and CAS typically display a stratigraphic progression upward to less negative values (Peng et al., 2011; Bao et al., 2012; Crockford et al., 2016, 2017) and later Ediacaran evaporites and CAS continue this trend (Fig. 7). This isotopic trend implies that a rapid transition occurred from anomalous conditions characteristic of the Marinoan glaciation (Bao et al., 2008; Cao and Bao, 2013), to conditions under a very different pO2-GPP-pCO2 combination than those occurring during other periods in the Proterozoic. Middle Ediacaran ∆17O values from evaporites for Siberian and Iranian sulfate evaporites as well as northern Australia CAS display less depleted ∆17O values, with the most negative values only reaching −0.29‰ (Fig. 7). Applying GPP calculations along with estimates of Ediacaran pO2 and pCO2 suggest high GPP levels during the Ediacaran Period relative to earlier in the Proterozoic. Although inferences about GPP may be hampered by isotopic resetting following sulfide-oxidation and dilution of sulfate before it is deposited in the sedimentary record, other isotopic insights from the sulfate record support an increase in GPP above typical low-mid Proterozoic levels (Fig. 9). The environmental conditions discussed above, together with the large Ediacaran sulfate evaporite deposits and growth in the marine sulfate reservoir, support high GPP and increased atmospheric and marine oxygen levels. Indeed, evidence of ≈16 mM sulfate has been suggested from fluid inclusions from the Ara Group of Oman (Brennan et al., 2004). Although Ediacaran ∆17O values are not as anomalous as post-Marinoan values, they are more negative than those of sulfate evaporites formed in modern environments (Fig. 7). Such values could be the consequence of higher pCO2 levels, possibly resulting from a slightly dimmer Ediacaran sun and reduced weathering capacity before the colonization of land plants (Berner, 2006, 1997). An alternative possibility is that lower GPP characterized the Ediacaran environment due to throttled phosphorous release from sediments under more oxic bottom-water conditions (Lenton et al., 2014; Sahoo et al., 2016; Stolper and Keller, 2018). While atmospheric and biospheric conditions recorded in ∆17O values tell a consistent story, sulfur isotope signatures bear a closer resemblance to complicated trace-metal records (Johnston et al., 2013; Kunzmann et al., 2015; Sahoo et al., 2016; Miller et al., 2017). Samples from Iranian and Oskoba Formation evaporites of Siberia span the range in δ34S recorded by post-Marinoan barites (Peng et al., 2011; Crockford et al., 2016, 2017) from values of ~20‰ in Iranian samples to ~45‰ in the Oskoba Formation (Fig. 7; Table 1). Such variation between sulfate evaporite deposits separated in time by ca. 15 Myrs, suggests a dynamic sulfur cycle over this interval with periods of intense microbial sulfur cycling and potential expansion and contraction of euxinic environments (Gomes and Johnston, 2017; Lang et al.,

2018), leaving an isotopically heavy marine reservoir, as observed in Oskoba Formation evaporite samples and previously reported from measurements of CAS (Fig. 7; Strauss et al., 2001; Johnston et al., 2005b; Goldberg et al., 2005; Hough et al., 2006; Gill et al., 2007; Tostevin et al., 2017). The Ediacaran δ18O record is similar to other isotopic records, with a large range of initial variation bracketed by more coherent trends that settle toward the Cambrian boundary (Fig. 7). While initial post-Marinoan values across multiple paleo-continents appear to be constrained to ~10 and 25‰, variation increases in ca. 600 Ma South China sections with extremely light values near 0‰ (Goldberg et al., 2005). This large range of variation has been interpreted as a result of both an increased proportion of water-oxygen incorporated into sulfate during sulfide oxidation that generated isotopically light values and enhanced euxinia driving oxygen isotope values heavier toward the PrecambrianCambrian boundary (Goldberg et al., 2005). Another possibility consistent with the expansion of euxinic environments is enhanced microbial sulfur cycling in sediments. However, ∆33S values appear to be similar to modern seawater values with averages of 0.022 and 0.015‰ for Iranian and Oskoba Formation samples, respectively. While enhanced dissimilatory sulfate reduction would result in more negative ∆33S values in marine sulfate, the onset of a significant reoxidative sulfur cycle at this time in the oceans with a rise to prominence of sulfur disproportionation (Kunzmann et al., 2017), may have counteracted this effect. In sum, the transition from the Proterozoic to the Phanerozoic revealed through Ediacaran sulfate isotope records appears to be complicated, but likely a consequence of the progressive oxygenation of the marine environment. While ∆17O results speak to intermediate conditions between typical Proterozoic and Phanerozoic states, further population of this record is needed to tie marine and terrestrial records to atmospheric conditions. Furthermore, changes in global temperatures (most notable through the Gaskiers glaciation) and dramatic shifts in the carbon isotope record were likely coeval with important changes to the Earth system, which the present evaporite record is unable to capture. The existing isotopic record of sulfate evaporites appears to be consistent with an increase in microbial sulfur cycling that progressively drove major isotopic values (δ18O and δ34S) of sulfate to become more positive. These trends are a predictable response to increased organic matter loading through enhanced primary productivity, which in turn would have driven ∆17O values closer to 0‰ than was likely possible at earlier times in Earth history. 5.7. Phanerozoic Phanerozoic sulfate isotope records (Claypool et al., 1980; Paytan et al., 1998, 2004; Masterson et al., 2016; Bao et al., 2008; Present et al., 2015) contrast sharply with the majority of the Proterozoic, having significantly less negative ∆17O values (< −0.32‰; Bao et al., 2008), coherent trends in δ34S and δ18O values, and slightly more positive ∆33S values (Fig. 7). These isotopic characteristics are a likely consequence of the progressive rise in atmospheric oxygen through the latest Proterozoic that appears to have continued into the Phanerozoic (Sperling et al., 2015; Stolper and Keller, 2018), potentially reaching modern levels with the colonization by land plants in the late Silurianearly Devonian (Lenton et al., 2016). Redox conditions in the marine environment coupled to a rise in atmospheric oxygen have largely been borne out through multiple lines of evidence, one of which being a growth in size of the marine sulfate reservoir, possibly a consequence of the advent of bioturbation that galvanized reoxidative sulfur cycling (Canfield and Farquhar, 2009; Tarhan et al., 2015). Here, we add another line of contrast between the Proterozoic and Phanerozoic by suggesting that an increase in atmospheric oxygen was tied to a rise in GPP (Fig. 9; Wing, 2013; Crockford et al., 2018). Higher Phanerozoic productivity together with lower pCO2 levels (Berner, 2006; von Paris et al., 2008; Franks et al., 2014; Dunlea et al., 2017) predict less 217

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negative ∆17O values, which is borne out in the isotopic record produced and compiled here. The impacts of this growth and diversification of the biosphere likely had consequences for the sulfur cycle that extended all the way to the mantle (Canil and Fellows, 2017). Such consequences include a growth in the size of the marine sulfate reservoir that is supported by largely coherent δ34S trends, albeit with some notable exceptions (Owens et al., 2013; Gomes et al., 2016). These coherent trends also suggest that geographically disparate records reflect a homogenous global reservoir. This observation, coupled to near uniformly positive ∆33S values, also speaks to an increase in microbial sulfur cycling across the Proterozoic – Phanerozoic transition (Fig. 7).

−0.44‰ provide evidence that the late Proterozoic environment was characterized by higher pCO2 levels than those during the Ediacaran or Phanerozoic times as well as lower GPP. High δ18O values potentially reflect a vigorous sulfur cycle leaving marine sulfate isotopically heavy. 6.4. Cryogenian (0.72–0.635 Ga) Highly negative ∆17O values within Marinoan aged strata indicate extremely high pCO2 levels toward the end and immediately after this Snowball Earth glaciation. Other aspects of the sulfur cycle remain poorly constrained due to incomplete isotopic records. Future work focused on the Cryogenian is needed to both better characterize the conditions immediately preceding each glaciation, the interglacial interval, and the dynamics of the sulfur and oxygen cycles in the subglacial oceans.

6. A speculative synthesis Here we have put forward a comprehensive isotopic record of Proterozoic sulfate from 313 samples from 33 different formations. This record, together with existing available data, confirms suggestions that atmospheric chemistry has significantly evolved through Earth's history and can be subdivided into three broad stages based on sulfate sulfur and oxygen multiple isotope records: S-MIF and no O-MIF (Archean), OMIF and no S-MIF (Proterozoic), and no S-MIF or O-MIF (Phanerozoic). We construct the first GPP curve across the Proterozoic (Fig. 9) based on atmospheric constraints coupled to the ∆17O record and argue that changes in primary productivity likely underlie many of the biogeochemical transformations that occurred during this eon. However, we emphasize that the GPP estimates put forward (Fig. 9) depend on existing estimates of Proterozoic atmospheric pCO2 and pO2, which vary widely and can thus provide very different productivity trajectories across the Proterozoic (Fig. 9). By providing the first empirical estimates of Proterozoic GPP, we highlight the biospheric consequences of evolving atmospheric compositions and beckon new data to reduce uncertainty in the record. Importantly, these changes in GPP appear to be borne out in other isotopic systems within sulfate and below we summarize highlights of our findings for the Proterozoic.

6.5. Ediacaran (0.635–0.542 Ga) The Ediacaran continues the trend from the mid-Proterozoic through the late Proterozoic of increased oxygenation of the surface environment, growth of the marine sulfate reservoir and growth in the size and complexity of the biosphere. Isotopic records suggest vigorous sulfur cycling that drove δ34S values of marine sulfate to positive values. Furthermore, ∆17O values suggest that the Ediacaran transitioned the Earth between Proterozoic and Phanerozoic surface environments with respect to the Productivity of the biosphere. Acknowledgements The authors thank Maya Gomes and Clint Scott for providing detailed and constructive reviews of this manuscript. Funding for this work was provided by the NSERC-CREATE CATP, NSERC PGS-D fellowship, McGill McGregor Fellowship, McGill Mobility and GREAT programs, Canadian Polar Continental Shelf Program, Northern Science Training Program, Mineralogical Association of Canada Foundation Scholarship and Travel Grant and the Agouron Geobiology Post-doctoral Fellowship Program. The Stable Isotope Laboratory at McGill University was supported by the FQNRT through the GEOTOP Research Center. GPH and BAW acknowledge support through the NSERC Discovery program. HB acknowledges funding from the strategic priority research program (B) of CAS (XDB18010104). AB acknowledges funding from the NSERC Discovery and Accelerator grants (RGPIN-316500).

6.1. Earliest Proterozoic (GOE) (2.5–2.0 Ga) Mirrored mass-independent isotopic trends in ∆17O and ∆33S across this interval suggest the establishment of an ozone layer immediately before the appearance of Earth's earliest evaporites. Further, a lack of any S-MIF in the evaporite record suggests that crustal S-MIF recycling did not have any significant impact on the sulfur isotope composition of seawater at the early stage of the GOE. If existing high pO2 estimates are correct, this oxygenation possibly coincided with the most productive biosphere in Earth history.

Appendix A. Supplementary data: Complete data compilation Supplementary data to this article can be found online at https:// doi.org/10.1016/j.chemgeo.2019.02.030.

6.2. Mid-Proterozoic (2.0–1.1 Ga)

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