Effects of a Transient Marine Sulfur Reservoir on Seawater Оґ18oso4

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Geochimica et Cosmochimica Acta 269 (2020) 257–269 www.elsevier.com/locate/gca

Eﬀects of a transient marine sulfur reservoir on seawater 18 d OSO4 during the Paleocene-Eocene Thermal Maximum

Weiqi Yao a,⇑, Adina Paytan b, Ulrich G. Wortmann a

a Department of Earth Sciences, University of Toronto, Toronto, Ontario M5S 3B1, Canada b Institute of Marine Science, University of California–Santa Cruz, Santa Cruz, CA 95064, USA

Received 1 March 2019; accepted in revised form 11 October 2019; Available online 25 October 2019

Abstract

Recent work suggests that during the Paleocene-Eocene Thermal Maximum (PETM) the volume of the oceanic oxygen minimum zone (OMZ) has expanded by one order of magnitude and acted as a transient reservoir for reduced sulfur. Fluc- 34 tuations in the seawater S-isotope ratios (d SSO4) can be used to estimate the size of this reservoir, but do not capture the gross fluxes associated with hydrogen sulfide reoxidation at the upper and lower boundaries of the OMZ. Isotope exchange processes during microbially mediated sulfate reduction as well as constant reoxidation of hydrogen sulfide to sulfate, how- 18 18 ever, affect the oxygen isotope ratio of seawater sulfate (d OSO4). Here we present a high-resolution (10-kyr) barite d OSO4 18 record. Our data shows that the d OSO4 value increased by 1.5‰ within 48 kyrs from the onset of the PETM and then returned to the pre-excursion value over the next 200 kyrs. Quantitative modeling suggests that the gross microbial sulfate reduction (MSR) flux was on the order of 4 1014 mol/yr, which is about one order of magnitude higher than the gross sulfate reduction flux in the modern ocean. Ó 2019 Elsevier Ltd. All rights reserved.

Keywords: Anoxia; Paleocene-Eocene Thermal Maximum; Seawater sulfate; Oxygen isotopes; Oxic sulfur cycle

1. INTRODUCTION reducing microbes; a process that has a preference for 32S and thus exerts a substantial influence on the sulfur isotopic 34 Seawater sulfate plays an important role in biogeochem- ratios (d SSO4) of the seawater sulfate pool (Kaplan et al., ical cycles. In organic-rich sediments, sulfate serves as the 1963; Brunner and Bernasconi, 2005; Wortmann et al., terminal electron acceptor to respire organic matter back 2001). As such, many studies have used the seawater 34 to carbon dioxide through microbial sulfate reduction d SSO4 to trace changes in the global sulfur cycle over 2 (MSR: 2CH2O+SO4 ? 2HCO3 +H2S; Berner, 1982; the Earth’s geological time (e.g., Claypool et al., 1980; Jørgensen, 1983). The partitioning of sulfur between the Paytan et al., 1998; Strauss et al., 2001; Newton et al., reduced and oxidized reservoirs exerts (together with car- 2004; Paytan et al., 2004). bon) the dominant control of oxygen in the ocean and Weathering on land oxidizes pyrite to sulfate which atmosphere (Walker, 1986; Wortmann and Chernyavsky, enters the ocean and is present in its dissolved form in pore 2007; Hurtgen, 2012). waters of sub-seafloor sediments. Once organic matter res- The transfer of sulfur from the oxidized reservoir to the piration has depleted all oxygen in the interstitial water in reduced reservoir is enzymatically catalyzed by sulfate- sediments, microbial sulfate reduction becomes the dominant respiration process with sulfide as the main metabolic end-product (Jørgensen, 1983; Reeburgh, 1983). In the ⇑ Corresponding author. presence of reactive iron, sulfide forms iron polysulfides E-mail address: [email protected] (W. Yao). and ultimately pyrite (Berner, 1982). However, a consider- https://doi.org/10.1016/j.gca.2019.10.019 0016-7037/Ó 2019 Elsevier Ltd. All rights reserved. 258 W. Yao et al. / Geochimica et Cosmochimica Acta 269 (2020) 257–269 able amount of sulfide is subsequently reoxidized when exposed to more oxic conditions either through mechanical mixing (bioturbation) or through diffusive loss (Jørgensen, 1982). This reoxidation flux is also known as the oxic sulfur cycle. Oxic sulfur cycling is of particular importance in bio- turbated sediments near the seawater-sediment interface, where the constant cycling between redox states has no net effect on the burial of reduced sulfur, but greatly affects 18 the oxygen isotope composition of marine sulfate (d OSO4; Turchyn and Schrag, 2006; Markovic et al., 2016). e e In the modern ocean, much of the oxic sulfur cycling Fig. 1. Simplified steps of microbial sulfate reduction. k and ex takes place in shelf areas where the anoxic part of the sed- denote the kinetic and equilibrium fractionation factors for iment column is shallow enough to overlap with the biotur- intracellular reduction and reoxidation respectively during MSR. Note that sulfite reduction and the release of sulfide can be bated zone. To a limited extent, oxic sulfur cycling can also reversible, but the backward flux is not well understood (Brunner occur in the water column of coastal upwelling zones and Bernasconi, 2005; Eckert et al., 2011; Holler et al., 2011). (Dugdale et al., 1977; Bru¨chert et al., 2003; Schunck Modified after Antler et al., 2017. et al., 2013). Recent work hypothesizes that this so-called cryptic sulfur cycling is more widespread and possibly 18 related to the reduction and oxidation of nitrogen ambient water d O(0‰; Van Stempvoort and Krouse, (Canfield et al., 2010b; Johnston et al., 2014). However, 1994); (b) Microbial sulfide oxidation which imparts a sulfate d18 ‰ d18 in the modern ocean, the maximum possible effect of cryptic OSO4 enrichment of 0-8 over water O(Van sulfur cycling on seawater sulfate oxygen is vanishingly Stempvoort and Krouse, 1994; Balci et al., 2012); (c) Micro- d18 small (Johnston et al., 2014). bial disproportionation where the sulfate OSO4 is 18 Recent studies suggest that the volume of oxygen-free increased by 8–21‰ over water d O depending on the pres- ocean water or the oxygen minimum zone (OMZ) has ence of oxidants (e.g., Fe(III), Mn(IV); Bo¨ttcher and increased substantially during the Paleocene-Eocene Ther- Thamdrup, 2001; Bo¨ttcher et al., 2001; Bo¨ttcher et al., 2005). mal Maximum (PETM; Dickens, 2000, 2011; Nicolo Which of these processes dominate depends on a variety et al., 2010; Dickson et al., 2012; Dickson, 2017; Yao of environmental conditions, chiefly variations in organic et al., 2018). As such, oxic sulfur cycling in the water col- matter availability and microbial activity (Turchyn and umn must have been more prominent. Here we present Schrag, 2006; Markovic et al., 2016). Previous studies sug- 18 high-resolution seawater d OSO4 data recorded in marine gest that oxic environments with limited amount of organic barite and use a numerical model to estimate the fluxes matter favor abiotic sulfide reoxidation whereas low- associated with the oxic sulfur cycle during the PETM. oxygen organic-rich environments favor microbially mediated sulfide oxidation (Jørgensen et al., 1991; Van 2. BACKGROUND Stempvoort and Krouse, 1994; Zopfi et al., 2001; Turchyn and Schrag, 2006; Markovic et al., 2016). Fig. 2 illustrates The microbially mediated reduction of sulfate to hydro- this schematically. gen sulfide progresses through a series of enzymatically cat- Since the reoxidation of sulfur proceeds with little or no alyzed reversible reactions (Rees, 1973; Brunner and S-isotope fractionation, the oxic sulfur cycle only affects the Bernasconi, 2005; Eckert et al., 2011; Wing and Havely, sulfate O-isotope ratio, but not the sulfate S-isotope ratio 2014; Antler et al., 2017). Each of these steps can affect oxy- (Turchyn and Schrag, 2006; Turchyn et al., 2010). The d34 gen isotopes of sulfate via kinetic and equilibrium exchange result is that the SSO4 records only the net flux of sulfate d18 reactions. The associated fractionation values vary between reduction while OSO4 will reflect the gross fluxes from 9‰ and 22‰ depending on the cell-specific sulfate reduction sulfate reduction and sulfide reoxidation. It is, however, rate (Fritz et al., 1989; Antler et al., 2017). Additionally, the important to understand that while the net sulfate reduc- reverse flux of intracellular sulfite to sulfate (and potentially tion flux might be small due to the reoxidation process, other reduced sulfur species) will add oxygen from cell- the total sulfate reduction flux can be much larger (See internal water which itself exchanges with the ambient seawa- Fig. 3). In the modern ocean, the gross flux of sulfate reduc- ter outside the cell (see Fig. 1). The resulting overall sulfate tion is estimated to be about ten times the net flux (Turchyn 18 d OSO4 equilibrium value is up to 29‰ higher than the and Schrag, 2004; Markovic et al., 2016). Here we will use 18 d18 d O of ambient seawater at 5 °C(0‰; Fritz et al., 1989; seawater OSO4 to shed light on the gross fluxes in and out Van Stempvoort and Krouse, 1994; Brunner and of the OMZ during the PETM. Bernasconi, 2005; Wortmann et al., 2007; Zeebe, 2010). In addition to the cell-internal effects during MSR, sulfate 3. METHODS O-isotope ratios are also affected by cell-external processes. Namely, some portion of the hydrogen sulfide produced dur- We use marine barite separated from sediment samples ing MSR is subsequently reoxidized to sulfate. Different from the Ocean Drilling Program (ODP) Leg 199 Hole 18 reoxidation pathways result in distinct d OSO4 signatures. 1221C Core 11X-3 spanning the PETM interval at approx- There are three known processes: (a) Abiotic sulfide oxida- imately 10-kyr resolution. The location of Hole 1221C dur- 18 tion which produces sulfate with d OSO4 values close to ing the PETM is within the Pacific equatorial upwelling W. Yao et al. / Geochimica et Cosmochimica Acta 269 (2020) 257–269 259

study show values around 0.70776 ± 0.00003 (2r), which compares favorably to the value of contemporaneous seawater (0.70770–0.70785; Hodell et al., 2007; Griffith et al., 2015). Based on the morphological examination and the Sr isotope-data we conclude that our samples are free of diagenetic and hydrothermal barite (Griffith et al., 2008; Griffith et al., 2015; Yao et al., 2018). Although some authors suggested that gas hydrate dissociation below the seafloor could lead to an increased supply of barium to the seafloor (e.g., Dickens et al., 2003), the absence of diagenetic barite in our cores and the constancy of the barite Sr/Ba ratio across the PETM interval (Paytan et al., 2007), suggest that this barium flux had no effect on our samples. Isotope ratios were obtained using a continuous flow isotope ratio mass spectrometer system (CF-IRMS, Finni- gan MAT 253 in continuous flow mode using a Conflo III open split interface) at the Geobiology Isotope Labora- tory at the Department of Earth Sciences (University of Toronto). Approximately 200 (±5%) lg barite samples Fig. 2. In the modern ocean (A), oxic sulfur cycling is limited to wrapped in a silver capsule are introduced into a Hekatech biotic processes in the topmost sediment layers in shelf sediments, high-temperature pyrolysis furnace at 1350 °C and pyro- and the much slower abiotic reoxidation of sulfide in deep-sea lyzed under helium atmosphere. Glassy carbon is used to sediments (e.g., Markovic et al., 2016). S-isotope ratios during the convert the liberated oxygen to carbon monoxide (CO) PETM suggest that, large parts of the ocean turned anoxic, or even gas for O-isotope measurement. The O-isotope values are sulfidic (Yao et al., 2018), resulting in increased cycling of reduced calibrated using four international standards with respect sulfur (B). to Vienna Standard Mean Ocean Water (VSMOW; Brand et al., 2009): USGS32 (25.4‰), IAEA-SO-5 (+12.13‰), NBS 127 (8.59‰), and IAEA-SO-6 (11.35‰). Repeated measurements of an in-house synthetic BaSO4 standard yield reproducibility of 0.29‰ (1r).

3.1. Statistical analysis

18 The d OSO4 value of seawater sulfate at any given time 18 (t) depends to a certain degree on the d OSO4 value of seawater sulfate at a given time before (t Dt), which allows us to apply a ‘‘local regression smoothing” technique 18 (LOESS; Cleveland, 1979) to estimate the d OSO4 value Fig. 3. Gross fluxes from sulfate reduction and sulfide reoxidation of seawater sulfate. We use the default LOESS module pro- 18 34 vided by the statistical software package R (R Core Team, control the seawater d OSO4 while the net flux controls the d SSO4. 2012). The 95% confidence interval is calculated for each data point from the standard errors returned by the LOESS zone (12.03°N, 143.68°W) at a palaeodepth of 3000 m function. 18 (Nunes and Norris, 2005; Winguth et al., 2012). We use To test the significance of the d OSO4 excursion data, the timescale of Ro¨hl et al. (2000) and Nunes and Norris we group the measurements into two groups, one ‘‘inside”’ (2005) to convert the depths of the associated sediments and the other ‘‘outside” the excursion. Fig. 4 shows a box to time relative to the onset of the PETM. Sedimentation plot showing the data distribution. A student t-test with rates are linearly interpolated between the known datum both groups shows that the resulting value of t (3.39) is lar- points. ger than the critical t-value (2.18) and that the p-value of Barite was extracted from the sediment using a sequen- 0.005 is significantly below the 0.05 threshold. We can thus tial leaching method (Markovic et al., 2016). All samples conclude that the means of the two groups are significantly have been scanned with a scanning electron microscope different at a 95% confidence level. (SEM) to test for the presence of diagenetic and hydrothermal barite which has considerably different morphology 3.2. Model setup and sizes from authigenic marine barite (Paytan et al., 2002; Griffith et al., 2008; Griffith et al., 2015). Addition- Here we expand upon the sulfur cycle model in Yao ally, we measured the strontium isotope ratio of our sam- et al. (2018) by including the additional terms for sulfate ples to ensure that the samples are of marine origin reduction and sulfide reoxidation. Fig. 5 shows the major 18 (Paytan et al., 1993, 2002). All of the samples used in this fluxes controlling d OSO4. These include fluxes from vol- 260 W. Yao et al. / Geochimica et Cosmochimica Acta 269 (2020) 257–269

OMZ and subsequently be reoxidized at the boundaries of the OMZ. The sulfide reoxidation flux is a function of the H2S concentration, the areal extent of the OMZ, and the vertical advection rate (see Eq. (5)). The background H2S concentration in seawater is zero outside the PETM interval and outside the OMZ. The estimated S-isotope fractionation associated with MSR in the OMZ is 58‰ (Eq. (7)), which is in line with fractionation factors observed in the environment of comparable geochemical settings to that of the PETM (e.g., 53‰ to 65‰ in Black Sea and various meromictic (stratified) euxinic lakes; Nakai and Jensen, 1964; Sweeney et al., 1980; Fry, 1986; Fry et al., 1991; Overmann et al., 1996; Ivanov et al., 2001; Canfield et al., 2010a; Zerkle et al., 2010). For the pre-PETM, we assume an ocean volume of 18 3 18 1.38 10 m with a seawater sulfate d OSO4 value of 10.8‰ (Claypool et al., 1980) and a marine sulfate concentration of 5 mM (Wortmann and Paytan, 2012). Fig. 4. Box plot showing the data distribution, the standard deviation, and the mean for samples inside and outside of the Note that the exact value of sulfate concentration used PETM peak. See Fig. 6 and Table 2 for how the samples are here is not critical. If sulfate concentrations were higher categorized. than assumed, the model calculated masses would be higher as well; if concentrations were lower, the results would be smaller. It would, however, not change the canic degassing, evaporite weathering and pyrite weather- mechanisms affecting the isotope values. The complete ing on land, evaporite burial, and MSR and reoxidation list of parameters is shown in Table 1. in sediments and seawater. We use the following assumptions: Using the above assumptions, we can calculate the mass of seawater sulfate as: Weathering and burial fluxes of evaporite remain con- M ðÞ¼t M ðÞþðt F ðÞþt F ðÞþt F ðÞAtðÞ stant throughout the model run since there is no sub- SO4 SO4 1 V WE WP F ðÞt F ðÞþAtðÞ F ðÞ stantial precipitation or dissolution of evaporites BE SED IN SED OUT t

during the PETM (Hay et al., 2006; Wortmann and F MSR INðÞþt F MSR OUTðÞþt F OX OMZðÞÞt Dt Paytan, 2012). ð1Þ MSR rate is a function of sulfate availability and organic matter quality and quantity (e.g., Canfield, 2001; where MSO4(t) denotes the mass of seawater sulfate [mol] at Wortmann and Chernyavsky, 2007). any given time step t; FV, FWE, and FBE denote the volcanic Although sea-level variations during the PETM are min- degassing flux, evaporite weathering flux, and evaporite or, we do consider the effect of these variations on shelf burial flux [mol/yr], which are assumed to be constant area and hence pyrite burial/oxidation (see Markovic throughout the model run; FWP(A) denotes the sulfate flux et al., 2015; Markovic et al., 2016). from pyrite weathering on land and continental shelves (see The globally averaged vertical advection in the ocean is details in Yao et al., 2018); FSED_IN(A) denotes the sulfate similar to the present at 4 m/yr (Broecker and Peng, influx for MSR in sediments as a function of shelf area; 1982). FSED_OUT denotes the sulfate reflux out of sediments; Since there is almost no reactive iron to bind the sulfide FMSR_IN and FMSR_OUT denote the forward sulfate flux into pyrite in seawater, the sulfide produced by MSR in for MSR and the backward flux of microbially exchanged the water column will accumulate in the core of the sulfate back to the OMZ; FOX_OMZ denotes the reoxidation

Fig. 5. Major ﬂuxes controlling the concentration and oxygen isotopes of seawater sulfate. Note that ﬂuxes from the OMZ are too small to measure outside the PETM interval (Jørgensen et al., 2004). See details in Table 1. W. Yao et al. / Geochimica et Cosmochimica Acta 269 (2020) 257–269 261

Table 1 Modelled sulfate fluxes and isotope ratios for the initial steady state. 34 18 Parameter Initial value d SSO4 [‰ d OSO4 [‰ Sources VCDT] VSMOW] 21 VSO4 = ocean volume 1.38 10 L [1] [SO4] = seawater sulfate concentration 5 mM 17.5 10.8 [2] 13 2 A1 = shelf area 3 10 m [3] 12 FWE = evaporite weathering flux 1 10 mol/yr 22 13 [4] 11 b FWP = pyrite weathering flux 6 10 mol/yr 20 0 [5] 11 FV = volcanic degassing flux 5 10 mol/yr 0 3 [6] 12 a a a FBE = evaporite burial flux 1.478205 10 mol/yr 17.5 10.8 [7] 11 a,b a FBP = pyrite burial flux 6.21795 10 mol/yr 25.51 [8] 13 a,b a FSED_IN = sulfate influx for MSR in sediments 5 10 mol/yr 10.8 Eq. (2); [9] 13 a a,c FSED_OUT = sulfate reflux out of sediments 4.9378 10 mol/yr 10.966 Eq. (3); [10] a FMSR_IN = forward sulfate flux for MSR within the 0 mol/yr 10.8 Eqs. (6) and OMZ (7) a,c FMSR_OUT = backward flux of microbially exchanged 0 mol/yr 14.1–9.8 Eqs. (6) and sulfate into the OMZ (7) c FOX_OMZ = sulfide reoxidation flux at the OMZ 0 mol/yr 0–21 Eqs. (4) and boundaries (5) R = reference ratio of O-isotope standard 0.0020052 [11] Dt = time step 100 years

Note: V = volume, A = area, F = flux, W = weathering, B = burial, V = volcanic, E = evaporite, P = pyrite, MSR = microbial sulfate reduction, OX = sulfide reoxidation, SED = sediments, OMZ = oxygen minimum zones (see details in the Methods section). a The initial steady state values calculated from other known parameters. b The fluxes affected by shelf area corresponding to sea-level variations. c The oxygen-isotope offset or equilibration between sulfate and ambient water (0‰). Here we list the actual numbers we used in the model, but in text we approximate the numbers with minimum significant figures. Reference sources: [1] Burke and Sengo¨r, 1988; [2] Paytan et al., 1998; Horita et al., 2002; [3] Yao et al., 2018; [4] Garrel and Lerman, 1984; Walker, 1986; Claypool et al., 1980; [5] Garrel and Lerman, 1984; Kump and Garrels, 1986; Krouse and Mayer, 2000; [6] Hansen and Wallman, 2003; Alt et al., 2010; [7] Garrel and Lerman, 1984; Walker, 1986; Claypool et al., 1980; [8] Yao et al., 2018; [9] Fritz et al., 1989; Turchyn and Schrag, 2004, 2006; Markovic et al., 2016; [10] Bo¨ttcher et al., 2005; Turchyn and Schrag, 2006; Markovic et al., 2016; [11] Assonov and Brenninkmeijer, 2003.

flux at the upper and lower boundaries of the OMZ (see M H2SðÞ¼t M H2SðÞþt 1 ðÞF MSR OMZðÞt F OX OMZðÞt Dt D Fig. 5); t denotes the model time step (100 years). ð4Þ The model first calculates the size of shelf area by using the sea-level estimates by Miller et al. (2005) (see Yao et al. where FMSR_OMZ denotes the MSR flux within the OMZ; (2018) for details). In the initial steady state, the sulfate MH2S(t) denotes the mass of H2S accumulated in the OMZ at any given time step. Note that H2S stands for reduction flux is the sum of 50% sea-level dependent MSR in shelf sediments and 50% MSR in abyssal sediments the sum of HS and H2S at seawater pH between 7 and 8 that are not affected by sea-level variations: (Penman et al., 2014). FMSR_OMZ and FOX_OMZ determine the net flux of H S. The maximum H S mass is consistent F ðÞ¼AtðÞ : F þ : F AtðÞ=A 2 2 SED IN 0 5 SED IN 0 5 SED IN 1 with the value obtained from the sulfur cycle model ð2Þ (8 1016 mol; Yao et al., 2018). The reoxidation flux out of the OMZ is a function of the where A(t) denotes shelf area at any given time step; A 1 H S concentration, the surface area of the OMZ, and the denotes the initial shelf area corresponding to the pre- 2 vertical advection rate, which is calculated as: PETM sea level (see Table 1; Yao et al., 2018); FSED_IN* denotes the steady-state MSR influx estimate correspond- F ðÞ¼t v ½H SðÞt A ð5Þ ing to the pre-PETM shelf area. OX OMZ 2 H2S The sulfate reflux out of sediments is calculated as the difference between the MSR influx and the pyrite burial where v denotes the globally averaged vertical advection flux: rate (4 m/yr; Broecker and Peng, 1982); [H2S] denotes the average H S concentration calculated from M and the F ðÞ¼t F ðÞAtðÞ F ðÞAtðÞ ðÞ 2 H2S SED OUT SED IN BP 3 volume of the OMZ (1.6 1017 m3; Yao et al., 2018); A denotes the surface area of the OMZ, which we calcu- where FBP(A) denotes the pyrite burial flux in deep sea and H2S 17 3 continental shelve sediments as derived from the sulfur late from the volume (1.6 10 m ; Yao et al., 2018) model in Yao et al. (2018) (see Fig. 5). assuming the average thickness of 750 m for the OMZ (Paulmier and Ruiz-Pino, 2009). The calculated A is The mass of H2S within the OMZ is determined by the H2S 14 2 sulfide input through MSR and the sulfide removal through about 60% of the global ocean area (3.61 10 m ; reoxidation: Jacobsen et al., 2000) (see Fig. 5). 262 W. Yao et al. / Geochimica et Cosmochimica Acta 269 (2020) 257–269

Table 2 Sample list of the oxygen-isotope data of marine barite. 18 Core, section, interval Depth [mbsf] Time relative to the onset d OSO4 [‰ VSMOW] SD (1r) of the PETM [kyrs] 1221C 11-3 0-3 153.40 232.90 11.01 0.29 1221C 11-3 5-8 153.45 215.60 10.73 0.29 1221C 11-3 16-19 153.56 177.40 11.39 0.29 1221C 11-3 20-23 153.60 163.50 10.87 0.29 1221C 11-3 25-28 153.65 146.20 10.90 0.29 1221C 11-3 30-33 153.70 128.80 11.17 0.29 1221C 11-3 35-38 153.75 111.50 11.25 0.29 1221C 11-3 40-43 * 153.80 94.10 11.42 0.29 1221C 11-3 45-48 * 153.85 76.80 11.03 0.29 1221C 11-3 50-54 * 153.91 55.95 11.56 0.29 1221C 11-3 54-58 * 153.95 48.35 12.29 0.29 1221C 11-3 58-62 * 153.99 42.80 11.98 0.29 1221C 11-3 62-66 * 154.03 37.26 11.52 0.29 1221C 11-3 66-70 * 154.07 31.74 11.48 0.29 1221C 11-3 70-72 * 154.10 27.60 11.91 0.29 1221C 11-3 76-78 * 154.16 19.36 10.60 0.29 1221C 11-3 78-80 * 154.18 16.60 10.94 0.29 1221C 11-3 80-82 * 154.20 13.84 11.37 0.29 1221C 11-3 84-86 154.24 8.32 10.83 0.29 1221C 11-3 86-88 154.26 5.56 10.76 0.29 1221C 11-3 88-90 154.28 2.80 10.84 0.29 1221C 11-3 90-92 154.30 0.00 10.92 0.29 1221C 11-3 92-94 154.32 2.80 11.04 0.29 1221C 11-3 94-96 154.34 5.56 10.92 0.29 1221C 11-3 96-98 154.36 8.30 10.56 0.29 1221C 11-3 98-100 154.38 11.10 10.88 0.29 1221C 11-3 105-108 154.45 20.70 10.98 0.29 1221C 11-3 110-113 154.50 27.60 10.83 0.29 1221C 11-3 140-142 154.80 69.10 11.06 0.29 18 Note: The d OSO4 data of marine barite are attained from ODP Leg199 Core 1211C-11–3. SD = 1-r standard deviation determined by repeated measurements of the in-house synthetic barium sulfate standard (99.9%). * 18 Denotes the samples inside of the PETM d OSO4 peak.

In addition to the fluxes which transfer sulfur in and out we use a similar approach and calculate the mass of each of the OMZ, we have to consider the oxygen-isotope isotope as: exchange fluxes during MSR (Fritz et al., 1989; Brunner 1000 M SO4 and Bernasconi, 2005; Wortmann et al., 2007; Turchyn M 16O ¼ ÀÁ ð8Þ d18O þ 1000 R þ 1000 et al., 2010). Since we know the net sulfate reduction flux, 18 we can calculate the forward and backward fluxes during ðd O þ 1000 M SO4 R M 18O ¼ ÀÁ ð9Þ sulfate reduction as a function of the expressed S-isotope d18O þ 1000 R þ 1000 fractionation factor assumed above (Rees, 1973; Brunner 16 and Bernasconi, 2005) as: where M16O and M18O denote the masses of sulfate OSO4 and 18O ; d18O denotes the isotopic delta ratio of oxygen F ðÞ¼t F ðÞþt F ðÞt ð Þ SO4 MSR IN MSR OMZ MSR OUT 6 in sulfate [‰]; R denotes the reference ratio of isotopic stan- 34 dard (0.0020052) for VSMOW (Assonov and a SMSR F ðtÞ¼ F ðtÞð7Þ Brenninkmeijer, 2003). MSR OUT a34S MSR IN MSR Note that certain fluxes (i.e., the evaporite burial flux 34 d18 where a SMSR* denotes the complete expression of S- and the forward MSR flux) changes with the OSO4 value isotope fractionation associated with MSR (70‰; Kaplan of seawater sulfate, which is only known after the summa- et al., 1963; Rudnicki et al., 2001; Wortmann et al., 2001; tion of all the fluxes. As such, we calculate these fluxes 34 Sim et al., 2011); a SMSR denotes the estimated S-isotope based on the delta value at the previous time step t-1 (see fractionation of 58‰ for MSR in the OMZ used in Eq. (10)). our model. These two values limit the ratio of the Sulfate reduction and sulfide reoxidation take place in backward flux to forward flux (i.e., FMSR_OUT/FMSR_IN) both the OMZ and marine sediments. For MSR within to be 83%. the OMZ, we assume that the forward sulfate flux (see 18 The above equations describe the general mass balance Fig. 1) has the same d OSO4 value as seawater sulfate, 18 in our model. In order to express the isotope mass balance, whereas the d OSO4 value of the backward sulfate flux W. Yao et al. / Geochimica et Cosmochimica Acta 269 (2020) 257–269 263

depends on the extent of the sulfate-water O-isotope M 16OðÞ¼t M 16OðÞþðt 1 F V 16OðÞþt F WE 16OðÞt 16 exchange during MSR. The respective sulfate OSO4 þ F ðÞAtðÞ F ðÞt ﬂuxes (i.e., F and F ) are calculated WP 16O BE 16O MSR_IN_16O MSR_OUT_16O F ðÞAtðÞ þ F ðÞt as: SED IN 16O SED OUT 16O

F MSR IN 16OðÞþt F MSR OUT 16OðÞt 1000 F ðtÞ F ðtÞ¼ÀÁMSR IN ð Þ þ F ðÞÞt Dt ð Þ MSR IN 16O 18 10 OX OMZ 16O 13 d OSO4ðt 1Þþ1000 R þ 1000 A similar equation applies to M18O. F ðtÞ F ðtÞ¼ÀÁ1000 MSR OUT According to the standard delta notion, we can calculate MSR OUT 16O 18 D OMSR OUT þ 1000 R þ 1000 the isotopic delta ratio of oxygen in seawater sulfate at any given time step t as: ð11Þ ÀÁ M ðtÞ=M ðtÞ R d18 18 18O 16O where OSO4(t-1) denotes the isotopic delta ratio of d OSO4ðtÞ¼ 1000‰ ð14Þ 18 R oxygen in seawater sulfate at time t-1; d OMSR_OUT denotes the O-isotope equilibration between ambient water (d18O=0‰) and cell-internal sulfite which returns 4. RESULTS into the marine sulfate pool via sulfate backward flux during MSR (see Fig. 1). Note that the oxygen Our results show that the oxygen isotopes of our marine isotope ratio of ambient water does not change through- barite samples vary between 10.5‰ and 12.3‰ across the out the model run and that all oxygen is derived from PETM with an average value of 11.1‰ (Fig. 6). The 18 it. d OSO4 value increases from about 10.8‰ to a peak value 16 The OSO4 flux of sulfate reoxidized from sulfide at the of 12.3‰ at 48 kyrs after the onset of the PETM and boundaries of the OMZ is determined by the reoxidation decreases gradually over the next 200 kyrs to its pre- flux and the associated O-isotope offset between sulfate excursion value. 18 and ambient water: The average value of our d OSO4 measurements is in 18 line with the published sulfate d OSO4 values (10.8‰) 1000 F ðtÞ recorded in evaporite gypsum in the Late Paleocene - Early F ðtÞ¼ÀÁOX OMZ ð12Þ OX OMZ 16O D18 þ R þ Eocene period (Claypool et al., 1980). However, similar to OOX OMZ 1000 1000 34 18 the d SSO4 signal (Yao et al., 2018), the d OSO4 data show a clear positive excursion across the PETM interval. d18 where OOX_OMZ denotes the steady O-isotope offset Although our data show considerable noise relative to the d18 ‰ between sulfate and ambient water ( O=0 ) caused signal, statistical analysis with a student t-test shows that by sulfide reoxidation at the boundaries of the OMZ. this excursion is statistically significant at a confidence Depending on the balance between abiotic and microbially interval of 95%, and we can thus reject the null hypothesis ‰ mediated sulfide reoxidation, this value ranges from 0 that both sample sets are statistically equal. Moreover, the ‰ (abiotic oxidation only) to up to 21 (microbially medi- temporal excursion over time (1.5‰) is significantly larger ated oxidation only) (Van Stempvoort and Krouse, 1994; than the analytical error (0.29‰,1-r). Bo¨ttcher et al., 2005; Balci et al., 2012). Similar to the processes for the OMZ, sulfate reduction 5. DISCUSSION in sub-seafloor sediments consumes sulfate and changes the sulfate S and O isotope ratios of the residual sulfate in the At typical ocean temperature and pH conditions, oxy- interstitial water pool. The S-isotope effect on seawater is gen isotopes of seawater sulfate do not equilibrate with described by a simple fractionation equation where seawa- ambient water oxygen within its residence time (0.5 Myrs; ter is treated a residual pool. However, the effect on seawa- Lloyd, 1968; Claypool et al., 1980; Chiba and Sakai, ter sulfate oxygen is more complex, and can be understood 1985; Jørgensen and Kasten, 2006). Therefore, the in terms of two processes: (1) Sulfide diffusing upwards to 18 d OSO4 value is thought to represent a record of seawater the sediment-water interface where it is reoxidized and 18 sulfate d OSO4 (Turchyn and Schrag, 2004, 2006; incorporates ambient water either microbially or abiotically Markovic et al., 2016). to sulfate; (2) The isotopic equilibration of seawater sulfate 18 The d OSO4 value is affected by a variety of processes: with interstitial water which has been modified by MSR. (a) Changes in the weathering rate of sulfur-bearing sedi- Unlike the OMZ, we cannot distinguish the O-isotope frac- ments; (b) Changes in the intensity of oxic sulfur cycling, tionation effects caused by the respective processes in sedi- e.g., changes in the size of shelf areas (Markovic et al., 18 ments. Rather, we assume a net O-enrichment for the 2016); and (c) Changes in the global sulfate reduction flux sulfate reflux out of sediments, which is treated as an offset which would affect the intensity of the microbially mediated d18 ‰ from ambient water O(0 ). Here we use an average O- O-isotope exchange between water and sulfate (Wortmann ‰ isotope offset of 11 , which is well within the limits of the et al., 2007). previously published values (7–15‰; Turchyn and Schrag, 18 Given the short duration of the PETM d OSO4 excur- 2004; Markovic et al., 2016). sion, changes in continental weathering are unlikely and Accordingly, the mass of sulfate 16O at any given 34 SO4 not corroborated by coeval changes in the d SSO4 signal. time step is expressed as: This is because the flux rates needed to create the observed 264 W. Yao et al. / Geochimica et Cosmochimica Acta 269 (2020) 257–269

(2018) thus proposed that the increased sulfate reduction rate is unrelated to processes at or within seafloor sediments but is likely to occur within the water column itself. Yao et al. (2018) hypothesized that increased ocean temperature and nutrient supply during the PETM caused the oxygen minimum zones to expand (Nicolo et al., 2010; Dickens, 2011; Dickson et al., 2012; Winguth et al., 2012). Once a critical pO2 threshold is reached (4 lM; Bru¨chert et al., 2003), sulfate reduction replaces oxic respiration as the main mode of organic matter remineralization in the water column. This has three important consequences: (A) The globally integrated microbial sulfate reduction rate increases, accelerating the oxygen isotopic exchange between sulfate and ambient seawater, driving the sulfate bound oxygen ratio towards more enriched values (Fritz et al., 1989; Van Stempvoort and Krouse, 1994; Wortmann et al., 2007); (B) In the absence of oxygen, hydrogen sulfide starts accumulating inside the OMZ, resulting in the buildup of a transient ocean internal reservoir of reduced sulfur; (C) Advective and diffusive processes will transport the reduced sulfur towards the boundaries of the OMZ, where the reduced sulfur is reoxidized to sulfate. Depending on the reoxidation pathway, this will either 18 enrich or deplete the d OSO4 value. We explore these processes by expanding the sulfur cycle model from Yao et al. (2018) to additionally include terms describing the transfer of oxygen atoms from water to sulfate. Similar to Yao et al. (2018), we assume an average sulfide concentration up to 0.5 mM within the OMZ (see Fig. 7D), well below the 1 mM threshold where chemocline upwelling and the subsequent release of H2S into the atmosphere become a possibility (Kump et al., 2005). These assumptions constrain the net sulfate reduction rate to 1.82 1012 mol/yr (Fig. 7C & E). The reoxidation flux at the OMZ boundaries is con- strained by the average sulfide concentration, the globally averaged vertical advection rate (4 m/yr), and the surface area of the OMZ (2.13 1014 m2; see Eq. (5) in the Meth- ods section). Since the average sulfide concentration reaches up to 0.5 mM, the reoxidation flux is limited to 4.27 1014 mol/yr at the height of the excursion (Fig. 7B). As such, conservation of mass requires a maxi- Fig. 6. (A) Oxygen isotope data of authigenic marine barite mum gross MSR flux of 4.28 1014 mol/yr (Fig. 7A), crystals across the PETM. (B) Sulfur isotope data of marine barite which is over 100 times the typical weathering fluxes. (Yao et al., 2018). Red circles and black circles denote the samples The seawater d18O is controlled by the balance ” ” SO4 groups ‘‘inside and ‘‘outside of the PETM excursions. Error bars between sulfide reoxidation and the microbially mediated are 1r. The gray envelopes denote the 95% confidence intervals of O-isotope exchange during MSR (see Fig. 3). Using the the LOESS regression. The shaded boxes indicate the main and recovery stages of the PETM (Nunes and Norris, 2005). fluxes above, we explore the parameter space of the model by considering the following two scenarios. In Scenario A, we assume that all the sulfide reoxidized at the OMZ 18 d OSO4 signal would require a 40-fold increase in evaporite boundaries is via abiotic oxidation contributing sulfate with 18 and pyrite weathering on land. Changes of this magnitude an average d OSO4 value of 0‰. To reproduce the 18 cannot be reconciled with the existing rock record nor with observed seawater sulfate d OSO4 excursion (Fig. 7F), 34 the d SSO4 record (Hay et al., 2006; Wortmann and our model requires that the isotope exchange flux during 18 Paytan, 2012; Yao et al., 2018). Similarly, the estimated MSR has a d OSO4 value of 14‰. In Scenario B, we 25-m sea-level rise during the PETM (Miller et al., 2005) assume that sulfide reoxidation at the OMZ boundaries would expand shelf areas by 4.3 1012 m2, adding less than proceeds via microbial disproportionation only and pro- 18 18 0.004‰ to the d OSO4 value. A massive increase in the glo- duces sulfate with an average d OSO4 value of 21‰. This 18 bal sulfate reduction rate is not without difficulty either, as requires the d OSO4 value of the isotope exchange flux dur- it would affect pyrite and organic matter burial. Yao et al. ing MSR to be around 10‰. Both scenarios suggest that W. Yao et al. / Geochimica et Cosmochimica Acta 269 (2020) 257–269 265

Fig. 7. Active oxic sulfur cycling across the PETM. (A) The MSR flux within the OMZ. (B) The sulfide reoxidation flux at boundaries of the OMZ. (C) The flux of hydrogen sulfide in and out of the OMZ. (D) Concentration of hydrogen sulfide within the OMZ. (E) Modeled sulfate 34 18 d SSO4 values versus our measurements (asterisk symbols). (F) Modeled sulfate d OSO4 values versus our measurements (asterisk symbols).

the average O-isotope fractionation between the micro- of the euphotic zone in the PETM oceans (1.2–9.4 bially exchanged sulfate and seawater during MSR is mol/(m2*yr); Ma et al., 2014). roughly halfway between the seawater d18O and the O- The above flux estimates are based on the assumption 18 isotope equilibrium of 29‰ (at 5 °C; Fritz et al., 1989). that the barite d OSO4 ratio records the average ocean sul- 18 However, in the absence of a good control of the magnitude fate d OSO4 ratio. There are two caveats here: (A) The of the O-isotope fractionation during MSR, we cannot dis- above fluxes are large enough to reduce the residence time tinguish which scenario is more likely to take place. of sulfate bound oxygen to about 1.6 104 years, which Ultimately, the gross sulfate reduction flux is limited by approaches the mixing rate time scales. As such, sulfate the amount of sulfate and the supply of organic matter oxygen ratios may not be uniform throughout the ocean; (OM). Since sulfate consumption is only affected by the (B) Most barite precipitation occurs below the photic zone much smaller net flux, OM supply is the only limiting fac- but before the supply of particulate organic matter has tor. The OM pool available is a function of the areal extent completely attenuated (i.e., a few hundred meters below of the sulfidic OMZ, however, some uncertainty exists with the sea surface; Martine-Ruiz et al., 2019). It is thus possi- respect to this value. Yao et al. (2018) proposed that about ble that the zone of barite formation and the zone of active 10% of the ocean water volume must have become sulfidic. sulfate reduction in the water column may overlap. Conse- Given 60% of the global ocean area have developed sulfidic quently, we cannot exclude the possibility that the observed 18 water with an average thickness of 750 m, a gross MSR flux sulfate d OSO4 signature is a mixture of barites which 14 2 of 4.28 10 mol/yr would consume 2 mol SO4/(m *year). record the condition above and inside the OMZ. Both A Reduction of one mole of sulfate requires two moles of and B would result in overestimating the magnitude of 2 18 organic carbon (OC). We thus need 4 mol of OC/(m *yr) the oceanic sulfate d OSO4 excursion. This will not change to sustain the modeled gross flux. This OC flux is within the fundamental processes described here but may result in the range of the estimated annual carbon export flux out smaller flux estimates. 266 W. Yao et al. / Geochimica et Cosmochimica Acta 269 (2020) 257–269

18 It is noteworthy that the d OSO4 signal takes a much two orders of magnitude. This paints a much more dynamic longer time to return to its pre-excursion value than the picture of the biogeochemical sulfur cycle, where the resi- 34 d SSO4 signal (Fig. 8). This is because the fast return of dence time of sulfur species in seawater can be much shorter 34 the d SSO4 value is facilitated by the transient storage of (less than 100 kyrs) than the millions of years estimate isotopically depleted sulfur in the OMZ. Once the stored based on external fluxes. This is even more pronounced sulfide returns to the ocean (as sulfate), the original mass for sulfate oxygen, where the microbial exchange fluxes balance between 32S and 34S is restored. No such transient during sulfate reduction exceed the gross flux of storage exists for the sulfate oxygen, which changes its iso- 4.28 1014 mol/yr by another order of magnitude, effec- topic ratio throughout the entire ocean. As such, the time- tively reducing the residence time to less than 10 kyrs. 18 scale for returning to its pre-excursion value is a function of Fluxes of this magnitude can change the d OSO4 ratio of the various sulfur fluxes into the water column (i.e., weath- smaller reservoirs like the oxygen minimum zone within ering, oxic sulfur cycling by bioturbation, etc.), which oper- years, and we can no longer assume that the ocean is well ate on much longer timescales than the buildup and release mixed with respect to sulfate during such events. Last but of H2S in the OMZ. not least, a transient sulfide reservoir like the oxygen minimum zone leaves almost no sedimentary record but does 6. SUMMARY affect the ocean sulfate S and O isotope ratios.

This study for the first time presents a high-resolution ACKNOWLEDGEMENTS 18 d OSO4 record of seawater sulfate across the PETM event. 18 We thank L. Kump and one anonymous reviewer for their com- The d OSO4 increases by 1.5‰ within 48 kyrs relative to the onset of the PETM and then decays to the pre- ments which greatly improved this manuscript. This work was sup- ported by a Discovery Grant of the Natural Sciences and excursion value over the next 200 kyrs. Quantitative analy- Engineering Research Council of Canada (NSERC) to U.G.W.; sis shows that this signal is consistent with the hypothesis the National Science Foundation (NSF) CAREER grant OCE- that large parts of the PETM ocean became sulfidic (Yao 0449732 to A.P.; and P. C. Finlay Q.C. President’s Fellowships et al., 2018). While the S-isotope data provides an estimate in Geology to W.Y. of the net transient transfer of sulfate to sulfide, the sulfate oxygen data provide limits on the total gross reduction and reoxidation fluxes. The gross flux estimates exceed the net REFERENCES fluxes by two orders of magnitude and are close to the upper limit of what the organic carbon export flux out of Alt J. C., Laverne C., Coggon R. M., Teagle D. A. H., Banerjee N. the euphotic zone can sustain. R., Morgan S., Smith-Duque C. E., Harris M. and Galli L. More importantly, however, we show that the fluxes of (2010) Subsurface structure of a submarine hydrothermal sulfur transformations within the oceanic water column system in ocean crust formed at the East Pacific Rise, ODP/ 11 can exceed the traditionally considered sulfur fluxes to IODP Site 1256. Geochem. Geophy. Geosy. , Q10010. Antler G., Turchyn A. V., Ono S., Sivan O. and Bosak T. (2017) and from seawater (i.e., weathering, pyrite burial, etc.) by Combined 34S, 33S and 18O isotope fractionations record different intracellular steps of microbial sulfate reduction. Geochim. Cosmochim. Acta 203, 364–380. Assonov S. S. and Brenninkmeijer C. A. M. (2003) On the 17O correction for CO2 mass spectrometric isotopic analysis. Rapid Commun. Mass Spectrom. 17, 1007–1016. Balci N., Mayer B., Shanks W. C. and Mandernack K. W. (2012) Oxygen and sulfur isotope systematics of sulfate produced during abiotic and bacterial oxidation of sphalerite and elemental sulfur. Geochim. Cosmochim. Acta 77, 335–351. Berner R. A. (1982) Burial of organic-carbon and pyrite sulfur in the modern ocean - its geochemical and environmental significance. Am. J. Sci. 282, 451–473. Bo¨ttcher M. E. and Thamdrup B. (2001) Anaerobic sulfide oxidation and stable isotope fractionation associated with

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