Oxygen isotope biogeochemistry of pore water sulfate in the deep ...

Geochimica et Cosmochimica Acta 71 (2007) 4221–4232www.elsevier.com/locate/gcaOxygen isotope biogeochemistry of pore water sulfate in thedeep biosphere: Dominance of isotope exchange reactions withambient water during microbial sulfate reduction (ODP Site 1130)Ulrich G. Wortmann a, *, Boris Chernyavsky a , Stefano M. Bernasconi b ,Benjamin Brunner c , Michael E. Böttcher d , Peter K. Swart ea Geobiology Isotope Laboratory, Department of Geology, University of Toronto, 22 Russellstr., Toronto, Ont., Canada M5S 3B1b ETH-Zurich, Geological Institute, 8092 Zurich, Switzerlandc Jet Propulsion Laboratory, California Institute of Technology, USAd Leibniz Institute for Baltic Sea Research, Seestr. 15, D-18119 Warnemünde, Germanye Marine Geology and Geophysics, Rosenstiel School of Marine and Atmospheric Sciences Miami, FL, USAReceived 2 October 2006; accepted in revised form 12 June 2007; available online 27 June 2007AbstractMicrobially mediated sulfate reduction affects the isotopic composition of dissolved and solid sulfur species in marine sediments.Experiments and field data show that the d 18 O SO4 2 composition is also modified in the presence of sulfate-reducingmicroorganisms. This has been attributed either to a kinetic isotope effect during the reduction of sulfate to sulfite, cell-internalexchange reactions between enzymatically-activated sulfate (APS), and/or sulfite with cytoplasmic water. The isotopic fingerprintof these processes may be further modified by the cell-external reoxidation of sulfide to elemental sulfur, and thesubsequent disproportionation to sulfide and sulfate or by the oxidation of sulfite to sulfate. Here we report d 18 O SO4 2 valuesfrom interstitial water samples of ODP Leg 182 (Site 1130) and provide the mathematical framework to describe the oxygenisotope fractionation of sulfate during microbial sulfate reduction. We show that a purely kinetic model is unable to explainour d 18 O SO4 2 data, and that the data are well explained by a model using oxygen isotope exchange reactions. We propose thatthe oxygen isotope exchange occurs between APS and cytoplasmic water, and/or between sulfite and adenosine monophosphate(AMP) during APS formation. Model calculations show that cell external reoxidation of reduced sulfur species wouldrequire up to 3000 mol/m 3 of an oxidant at ODP Site 1130, which is incompatible with the sediment geochemical data. Inaddition, we show that the volumetric fluxes required to explain the observed d 18 O SO4 2 data are on average 14 times higherthan the volumetric sulfate reduction rates (SRR) obtained from inverse modeling of the porewater data. The ratio betweenthe gross sulfate flux into the microbes and the net sulfate flux through the microbes is depth invariant, and independent ofsulfide concentrations. This suggests that both fluxes are controlled by cell density and that cell-specific sulfate reduction ratesremain constant with depth.Ó 2007 Elsevier Ltd. All rights reserved.1. INTRODUCTION* Corresponding author.E-mail address: uli.wortmann@utoronto.ca (U.G. Wortmann).Microbial sulfate reduction is the major pathway of organicmatter (OM) oxidation in coastal-marine and continental-shelfsediments (Jørgensen, 1982), and is afundamental process linking the geochemical cycles of carbon,sulfur, and oxygen (e.g., Berner, 1982; Schidlowskiet al., 1983; Garrels and Lerman, 1984; Wortmann andChernyavsky, 2007). Sulfate-reducing microorganisms reduceSO 42according to the following net reaction:SO 42þ 2CH 2 O ! H 2 S þ 2HCO 3ð1Þ0016-7037/$ - see front matter Ó 2007 Elsevier Ltd. All rights reserved.doi:10.1016/j.gca.2007.06.033

4222 U.G. Wortmann et al. / Geochimica et Cosmochimica Acta 71 (2007) 4221–4232Although several details of the fractionation process remaincontroversial, the overall process is well understoodand can be described as the sum of several mass dependentfractionations during the stepwise reduction of sulfate tosulfide (Fig. 1) and the ratio between the forward and backwardreactions (Rees, 1973; Brüchert, 2004; Brunner andBernasconi, 2005). Culture experiments with dissimilatorysulfate reducers and field data show that the evolution ofthe d 18 O SO4 2 value during progressive sulfate reductionEq. (1) is dependent on the oxygen isotope compositionof the water (e.g. Mizutani and Rafter, 1973; Fritz et al.,1989; Böttcher et al., 1998, 1999; Brunner et al., 2005).Thus, if water is strongly depleted in 18 O compared to sulfate,the oxygen isotope ratio of sulfate will decrease whileits sulfur isotope composition increases (e.g. Mizutani andRafter, 1973; Fritz et al., 1989; Brunner et al., 2005). However,the limited number of studies conducted so far do notagree on the cause of this isotope effect and several modelshave been proposed:(A) Microcosm experiments have conclusively demonstratedthat isotope exchange reactions between sulfateand water are the dominant control factor ofthe d 18 O SO4 2 value (e.g., Mizutani and Rafter,1973; Fritz et al., 1989; Böttcher et al., 1998; Brunneret al., 2005; Knöller et al., 2006). However, the possibilityof a kinetic d 18 O SO4 2 fractionation is still discussedin the literature describing marineenvironments. This interpretation is based on theobservation that in certain studies the measuredd 18 O SO4 2 and d 34 S values of dissolved sulfate showa linear correlation (e.g. Aharon and Fu, 2000; Bottrellet al., 2000; Mandernack et al., 2003). Thereported ratios between the fractionation factors ofO and S vary from 1:1.4 to 1:4 (Mizutani and Rafter,1969; Aharon and Fu, 2000). We will thereforeexplore whether this hypothesis is a good explanationfor the ODP Site 1130 data.(B) The d 18 O SO4 2 is a function of an oxygen isotopeexchange between metabolic intermediates and cytoplasmicwater during microbially-mediated sulfatereduction. Fritz et al. (1989) observed that the d 34 Sand d 18 O SO4 2 values of aqueous sulfate in a groundwater environment initially increased together, butthat the d 18 O SO4 2 value asymptotically approacheda constant value whereas the d 34 S value continuedto increase. This observation agrees with results fromanoxic pore waters of marine sediments (Zak et al.,1980; Böttcher et al., 1998, 1999, 2001). It was shownexperimentally that the final steady state d 18 O SO4 2value depends on the d 18 O value of the ambient water(Mizutani and Rafter, 1973; Fritz et al., 1989; Brunneret al., 2005) suggesting an isotope exchange reactionbetween ambient water and sulfate. Theexperimentally determined steady state value for bacterialcultures (29‰ at 5 °C, Fritz et al., 1989) issomewhat lower than the equilibrium value predictedfrom high temperature experiments (36.4‰ and33.6‰, Lloyd, 1968; Mizutani and Rafter, 1973,respectively). As oxygen isotope exchange reactionsbetween water and sulfate are extremely slow atambient temperature and circumneutral pH (Zaket al., 1980; Chiba and Sakai, 1985), it has been suggestedthat this exchange must take place betweenenzymatically activated sulfate (adenosine phosphosulfate,APS) or sulfite and cytoplasmic water (Mizutaniand Rafter, 1973; Fritz et al., 1989).(C) The d 18 O SO4 2 value is a function of (A and B). So far,this possibility has not been studied in great detail,but only mentioned (Fritz et al., 1989; Brunneret al., 2005). The combined effect of a kinetic isotopefractionation and an oxygen isotope exchange withambient water would lead to an offset of the equilibriumisotope value and create an apparent equilibriumisotope factor, which would be greater thanthe actual equilibrium factor.Cytoplasmic Membranef 1 f 2 f 3 f 4 f 5SO 2— 4 SO 2— 4 APSSO 2— 3 H 2 Sb 1 b 2 b 3 b 4 b 5H 2 SCytoplasmAMPH 2 OH 2 OFig. 1. The major fractionation steps and associated fluxes during microbial reduction of sulfate. The overall isotope effect depends on thesum of the individual steps and the ratio between the forward and backward fluxes. Modified after Brunner and Bernasconi (2005).

Oxygen isotope biogeochemistry of pore water sulfate... 4223(D) d 18 O SO4 2 isotope effects are mainly a product of theoxidative part of the sulfur cycle where H 2 S is oxidizedto S 0 2and then disproportionated to SO 4and sulfide. The net reaction describing this processcan be written as4H 2 O þ S 0 2! 3H 2 S þ SO 4þ 2H þð2ÞBöttcher et al. (2001, 2005) observed that the abovereaction results in a net isotope effect for d 18 O SO4 2 ofup to 21‰ at 28–35 °C. They suggested that this isotopeeffect may be facilitated via the formation of sulfite as ametabolic intermediate, which would allow for the oxygenisotope exchange with ambient water. This processhas been used to explain the measurements from interstitialwaters (e.g. Blake et al., 2006; Turchyn et al.,2006) and to explain the d 18 O SO4 2 value of ocean water(Turchyn and Schrag, 2006).(E) Oxidation of sulfide to sulfate with water or dissolvedoxygen. For example, a combination of hypothesis(A and D) was suggested by Ku et al. (1999) andLu et al. (2001).In the following, we present d 18 O SO4 2 data for dissolved2SO 4from interstitial water samples of ODP Leg 182(Feary et al., 2000a), and investigate which of the abovehypothesis best explains the Site 1130 data.1.1. Geologic backgroundThe Great Australian Bight (GAB) forms the centralembayment of Australia’s southern continental margin, locatedbetween 124°E and 134°E and 32°S and 34°S. It representsthe largest cool water carbonate depositional realmon Earth today (Feary and James, 1998) and was the focusof ODP Leg 182 (Feary et al., 2000a). Although the coolwatercarbonate ramp of the GAB represents an unusualsystem today, it is considered a modern analogue of theMesozoic carbonate ramps (Feary and James, 1998) thathost a large share of the world’s petroleum resources.Ocean Drilling Program Leg 182 drilled two transectsthrough this carbonate ramp (see Fig. 2) and recovered244 interstitial water samples down to 500 m below seafloor(mbsf). The interstitial water profiles of these cores indicatethat the margin contains a complex system of differentbrines, with salinity values up to three times that of seawater(see Fig. 3, and Hine et al., 1999; Feary et al., 2000a;Swart et al., 2000; Wortmann, 2006). These brines mayhave been generated during sealevel lowstands on the largeadjacent shelf and emplaced into the sediments of the uppercontinental slope under the influence of a hydraulic head(Swart et al., 2000; Jones et al., 2002). The geochemical signatureof the brine corroborates a seawater origin, andshows that the brine carries up to 84 mM of sulfate. AtODP Site 1130, the advecting brine results in the supplyof SO 42from below, where SO 42concentrations increasewith depth until they stabilize at 300 mbsf at 65 mM (Fearyet al., 2000a). Modeling the conservative chloride ion concentrationdata of Site 1130 shows that Site 1130 is influencedby strong upward advective flow on the order of2 mm/year (Wortmann, 2006). Inverse reaction-transportmodeling of the sulfate concentration data of this site, suggeststhat volumetric sulfate reduction rates (SRR) vary between600 and 65 pmol/cm 3 /year (Wortmann, 2006). Asthese carbonate sediments contain only small amounts ofiron, hypersulfidic conditions prevail throughout most ofthe cores recovered during ODP Leg 182 (Feary et al.,2000a). The sulfur geochemistry of these sites is unusualas the isotopic difference of dissolved sulfide and sulfate approaches70‰ (see Fig. 4, and for a detailed discussion seeWortmann et al. (2001) and Wortmann (2006)).33 S127 E 128 E129 E100200Eyre1134Terrace100011321130Jerboa-11126112911271131113350034 S2000300011284000GAB-13B(alternate)45000 50 kmFig. 2. Locations of the sites drilled during ODP-Leg 182. Contour lines are in m and show water depth. Figure taken from Feary et al.(2000a).

4224 U.G. Wortmann et al. / Geochimica et Cosmochimica Acta 71 (2007) 4221–4232567 mmol/l Cl - 567 mmol/l Cl -1300 mmol/l Cl - 1300 mmol/l Cl -Fig. 3. Isohaline surfaces derived from interpolation of the Cl data of Sites 1130 and 1132. The fact that the Cl concentrations stabilize at asub-horizontal level, suggests that the brine is constantly replenished at depth, possibly by lateral advection. The horizontal distance betweenthe two sites is about 11.6 km, seismic data courtesy of D.A. Feary.Depth [m]-50 0 500SO 4 [mM] DataH 2 S [mM] Data50δ 34 S SO 4 [ 0 / 00 VCDT] Dataδ 34 S H 2 S[ 0 / 00 VCDT] Data1001502002500 / 00[VCDT]SO 4 [mM] ModelH 2 S [mM] Modelδ 34 S SO 4 [ 0 / 00 VCDT] Modelδ 34 S H 2 S [ 0 / 00 VCDT] Model300-50 0 50[mM]Fig. 4. Wortmann et al. (2001) reported d 34 S data from ODP Site 1130 which showed an isotopic difference of more than 70‰ between coexistingdissolved H 2 S and SO 24. They argued that under hypersulfidic conditions and in the absence of any oxidants, this difference must becaused by unusually high S-fractionation factors during microbial sulfate reduction. Reaction transport modeling of this system suggests thatthe fractionation factor was at least 65‰ (Wortmann et al., 2001). A theoretical framework how to explain these large fractionation factorswas given by Brunner and Bernasconi (2005). Note that the modeling results presented in this figure have been obtained by the same nonsteadystate parametrization used throughout in this paper, and that Fig. 7 demonstrates that the observed d 34 S-signatures are not an artifactof the 1-D modeling approach.However, the interstitial water data of the upper 30 mbsfat ODP Site 182 are difficult to explain within the frameworkof a steady state diagenetic model. Previous publications(Wortmann et al., 2001; Wortmann, 2006) suggestedthat the upper 30 mbsf might be affected by an unknownmixing process. While it is difficult to envision a physicalprocess which provides for a steady state mixing depth of30 mbsf, non-steady state mixing by sedimentation eventsis a plausible process. Huuse and Feary (2005) suggestedthat the youngest sediments in the Great Australian Bightsediments might be contourites. Thus we assume that thesediments at Site 1130 are redeposited older sediments,which provides a plausible mechanism for seawater irrigation,and allows us for the first time, to successfully modelthe upper 30 mbsf of OPD Site 1130 (see Fig. 5).2. METHODSThe interstitial water samples were taken on board theJOIDES Resolution following the procedures given in

Oxygen isotope biogeochemistry of pore water sulfate... 4225Site 1130 14 C dataSite 1130 new age modelSite 1130 old age modelCl -New age modelOld age modelmbsf01020300 cm/y3.12 cm/y0 cm/y2.86 cm/y26 cm/kymbsf [m]010203040405010 0 10 1 10 2 10 3 10 4 10 5Age [yrs B.P.]50500 1.000 1.500Cl - [mM]Fig. 5. Comparison between the Site 1130 14 C age data, and the age model used in our model. We assume that the sediments at Site 1130 arereworked drift sediments (Huuse and Feary, 2005) and have been redeposited about 20–30 ky after their initial deposition. Note that themodel assumes two major re-sedimentation events (0.2–0.5 ky B.P. and 10.5–10.7 ky B.P.) and two prominent hiati (0–0.2 ky B.P, and 0.5–10.5 ky B.P.). The event and hiati timing and duration have been obtained from inverse modeling of the conservative Cl data of Site 1130,but are otherwise unconstrained. 14 C age data from Hine et al. (2002).Feary et al. (2000b). Samples were treated immediatelyafter collection by adding 100 lL of a saturated zinc acetatesolution to 5 mL of sample, to precipitate all H 2 S and inhibitfurther activity of sulfate-reducing microorganisms. Theprecipitated ZnS was separated using a centrifuge, and thesupernatant was filtered through a 0.45 lm membrane filterand acidified with HCl. The sulfate was precipitated asBaSO 4 by the addition of BaCl 2 within one hour after acidification.Samples were centrifuged, washed with hot water,and dried. For d 18 O SO4 2 measurements, about 135 lgBaSO 4 was added to Ag cups and pyrolyzed at 1400 °Con a Hekatech HT-EA using helium as a carrier gas. Theproduced CO was routed through an Ascarite trap, separatedon a Molsieve 5A, and subsequently measured on aThermo Finnigan Mat 253 in continuous flow mode usingthe Conflo III open split interface. The system was calibratedby using the international standards USGS 32(25.7‰ V-SMOW) and NBS 127 (9.3‰ V-SMOW).IAEA-SO-6 was also measured ( 11.34‰ V-SMOW) butnot used for calibration purposes since all data falls in therange between NBS 127 and USGS 32. Analytical reproducibilityof the measurements was determined by runningseveral replicates of an in-house standard (BaSO 4 withd 18 O = 12.38‰) with each run. We report the accumulatederror (1r) of in house standards measured at the beginning,middle and end of each run from 12 individual runs (i.e., 3per run = 36 total) as ±0.26‰. The data are reported in theconventional delta notation with respect to V-SMOW.Oxygen isotopic measurements of H 2 O from interstitialwater samples were made in the Division of Marine Geologyand Geophysics at the University of Miami using awater equilibration system (WEST) attached to an EuropaGEO. This setup determines the oxygen isotopic compositionfrom CO 2 which has been injected into serum bottles(5 cm 3 ) at slightly above atmospheric pressure containing0.5 cm 3 of sample. The samples are subsequently equilibratedat 40 °C for 12 h without shaking. The process isentirely automated with the CO 2 being injected and retrievedusing an autosampler and the gas being transferredto a dual-inlet mass spectrometer through a cryogenic trap(at 70 °C) to remove water. The precision of this methodfor oxygen, determined by measuring 59 samples of ourinternal standard, is ±0.07‰. All data are reported in‰ relative to V-SMOW using the conventional deltanotation.The equilibrium values reported here are based on theexperimental steady state values reported by Fritz et al.(1989), and we use the following relation to calculate theequilibrium values reported in Fig. 8.d 18 O SO4 2 ¼ 29:772 T 0:1599 ð3Þwhere T denotes the shipboard measured temperature profileof OPD Site 1130 (Feary et al., 2000a), which weapproximated asT ¼ 10:5 þ z 0:037ð4Þwhere z denotes the depth in meters below seafloor.Total iron of selected samples was measured by X-rayfluorescence spectroscopy (Philips PW 2400, equipped withRh-tube) on fused borate glass beads with a precision (2SD)of 1.1% (Dellwig et al., 2002). Total inorganic reducible sulphur(TRIS) which is considered to essentially representpyrite sulfur, was extracted from the sediments accordingto the one-step hot Cr(II)Cl 2 distillation method (Fossingand Jørgensen, 1989) where the H 2 S was trapped quantitativelyas Ag 2 S in an AgNO 3 solution and quantified

4226 U.G. Wortmann et al. / Geochimica et Cosmochimica Acta 71 (2007) 4221–4232gravimetrically (Böttcher et al., 2004). Pyrite-iron contentswere calculated from TRIS contents using the ideal stoichiometryof pyrite. Iron, leachable with buffered Na-dithionitesolution (Canfield, 1989) was also measured on twosamples (1130A-7-H3 and 1130A-21X-3) yielding 0.05 and0.04 dry wt%, respectively.2.1. Reaction transport model formulationThe distribution of dissolved interstitial water species affectedby diagenetic or biologically mediated conversionprocesses can be described as a process which involvestransport by diffusion, transport by advection, and a consumptionor production function (Berner, 1964, 1980;Boudreau and Westrich, 1984; Boudreau, 1996). The standarddiagenetic equation (Berner, 1980) for a dissolved speciesdescribes the change of concentration with depth andtime as a function of diffusion, advection, porosity, andconsumption or production:oðuCÞot ¼zhoðuCÞo D B ozþ uðD i þ DÞ oCozozioðuxCÞozwhere: C is the concentration, D are diffusion coefficients todescribe diffusion due to bioturbation (D B ), diffusion due toirrigation (D i ) , and the molecular diffusion term D. Thesymbol u stands for the porosity, the advection velocity isdenoted x, time is expressed as t, and the reaction or productionfunction is termed f.In the following, we will assume that there is no bioturbation,no diffusive irrigation, and that porosity changesdowncore only as a result of compaction (for a more detaileddiscussion see, Berner, 1980; Wortmann, 2006; Chernyavskyand Wortmann, 2007). Thus we can simplify Eq.(5) and writeu oCot ¼ u oC oDoz oz þ uD o2 Coz þ D ou oC2 oz ozux oCozThese equations can now be used for inverse modelingto quantify the advection velocity, and the net volumetricSRR (f) as a function of depth (see e.g., Berg et al.,1998). The diffusion coefficient used for SO 42is computedas a function of shipboard measured temperature andporosity using the parameters given by Boudreau (1996).The reduction function describing the sulfate reduction isalmost identical to the one given by Wortmann (2006),but unlike the model used by Wortmann (2006) and Wortmannet al. (2001), here we include two major sedimentationevents in the upper 15 mbsf (Hine et al., 2002). Thesedimentation event was modeled as a depositional eventof 5.72 m of sediment within 200 years 10 ky ago, and a secondevent, depositing 9.36 m of sediment within 300 yearsabout 200 years ago. These changes increase the estimatedupwelling velocity twofold to 1.2 · 10 10 m/s, and allow usto successfully model the chemical profiles in the upper30 mbsf, which was not possible previously (see e.g., Wortmannet al., 2001; Wortmann, 2006). All modeling wasdone with REMAP (Chernyavsky and Wortmann, 2007)which was modified to include a module to calculate isotopeexchange reactions.ufufð5Þð6Þ2.2. Modeling the kinetic oxygen isotope effects in sulfateIn order to model kinetic oxygen isotope effects in sulfatewe assume that this effect is proportional to the volumetricnet reduction rate f. We thus express the isotopeeffect similar to the approach of Jørgensen (1979), and writef ðS 16 O 4 Þ¼a½S 16 O 4 Š½SO 4 Šþða 1Þ½S 16 O 4 Š f ð7Þf ðS 18 ½S 18 O 4 ŠO 4 Þ¼a½SO 4 Š ða 1Þ½S 18 O 4 Š f ð8Þwhere values in square brackets denote concentrations.Note, that S 18 O is merely a notation to describe the concentrationof 18 O in a sulfate molecule but in no way impliesthe existence of a molecule with 4 18 O isotopes (whichwould be extremely rare under natural conditions). CombiningEqs. (5) and (7), we obtainu o½S16 O 4 Šot¼u o½S16 O 4 Š oDozux o½S16 O 4 Šozoz þ ½S 16 O 4 ŠuDo2 þ D ou o½S 16 O 4 Šoz 2 oz oza½S 16 O 4 Šu½SO 4 Šþða 1Þ½S 16 O 4 Š f ð9Þand similarly for 18 O. This allows us to describe the evolutionof the d 18 O SO4 2 values as a function of depth, a and f.2.3. Modeling the oxygen isotope exchange between waterand sulfateIn the following section, we do not distinguish whetherexchange reactions occur within the cell or outside the cell.We furthermore assume that there is no kinetic oxygen isotopefractionation during microbial sulfate reduction, i.e.,that the d 18 O SO4 2 value depends only on the exchange reactionbetween water and SO 2 4.As equilibration reactions require zero net flow, we candescribe the oxygen isotope effect of the exchange reactionon the pore-water sulfate pool in terms of a closed loop.There, dissolved sulfate leaves the pore-water pool and entersa box where its oxygen isotope signature is modified bythe exchange reaction, and afterwards returns back into thepore-water sulfate pool (Fig. 6). As before, we will only describethe volumetric exchange rates. Thus, the evolution ofthe isotopic composition of the pore-water sulfate reservoircan be described in terms of the volumetric exchange flux b,the isotopic enrichment during the exchange process and afactor k describing the extent of the exchange reaction.Since we assume that dissimilatory sulfate reduction hasno effect on the d 18 O SO4 2 value, the flux of 16 O and 18 Oisproportional to the initial ratio of the given isotope to thetotal sulfate amountf ðS 16 O 4 Þ¼f ½S16 O 4 Šð10Þ½SO 4 Šand similarly for f( 18 O).We assume that there is no isotope effect on thed 18 O SO4 2 composition for the flux associated with the exchangereaction and we can write this flux as being proportionalto the initial ratio of the given isotope to the totalsulfate amount as in Eq. (10). Since the total concentration

Oxygen isotope biogeochemistry of pore water sulfate... 4227sulfate reducing bacteriasulfate reduction {f } • δ 18 O SO4 porewaterSulfate inporewaterδ 18 O SO4 cytoplasm {δ Out }=δ 18 O SO4 porewater {δ In } • (1-k) +k•(δ H 2 O + ε)exchange flux {b} • δ 18 O SO4 porewaterexchange flux {b} • δ 18 O SO4 cytoplasmFig. 6. The principal fluxes involved in oxygen isotope exchange reactions between sulfate and water. Note that our model makes noassumption whether the exchange reactions happen cell-internal or cell-external.2of SO 4is not changed by the exchange reaction, we havethe additional constraint that the sum of the fluxes of 18 Oand 16 O cannot add any new sulfate, i.e., they must cancelbðS 16 O 4 ÞþbðS 18 O 4 Þ0ð11Þwhere b denotes the exchange velocity. Thus the isotopiccomposition of the output flux becomes simply a functionof the isotope equilibrium value and we can write it as acombination of the input and output flux asbðS 16 ½S 16 O 4 Š1O 4 Þ¼b þð12Þ½SO 4 Š 1 þðd out =1000 þ 1ÞR 0bðS 18 ½S 18 O 4 ŠO 4 Þ¼b þ ðd out=1000 þ 1ÞR 0ð13Þ½SO 4 Š 1 þðd out =1000 þ 1ÞR 0where R 0 refers to the isotopic ratio of V-SMOW, andd out ¼ d in ð1 kÞþkðd H2 O þ Þ ð14Þand d H2 O denotes the oxygen isotopic composition of thepore-water H 2 O. The reaction transport model to describethe oxygen isotope exchange for 16 O is then a combinationof Eqs. (6, 10 and 11)u o½S16 O 4 Šot¼u o½S16 O 4 Š oDozux o½S16 O 4 Šozand similarly for [S 18 O 4 ].oz þ ½S 16 O 4 ŠuDo2 oz 22.4. Modeling cell external oxidationþ D ou o½S 16 O 4 Šoz ozuf ðS 16 O 4 ÞþubðS 16 O 4 Þð15ÞCell external oxidation of H 2 S can be achieved througha solid oxidant like ferric oxihydroxide. As bioturbationwill only affect the benthic boundary layer, we can treatour oxidant as a solid, and writeoðð1uÞ½X s ŠÞot¼ oðð1 uÞx s½X s ŠÞozð1 uÞf n ð16Þwhere X s denotes a solid oxidant, f the gross flux of SO 4 2 ,x s the sedimentation rate, and n the molar ratio of the oxidationreaction.3. RESULTS AND DISCUSSIONUnder steady state conditions with no lateral changes ofboundary conditions, results of a 1-D model must equal theresults of a 2-D model. However, if lateral changes (e.g., ofsulfate concentration or biological activity) do occur, 1-Dmodels may somewhat overestimate sulfate reduction rates.At present, the hydrogeology of ODP-Site 1130 is not sufficientlyconstrained to allow for a detailed 2-D model.However, we can set up a 2-D model which is sufficientlysimilar (i.e., it uses the same spatial scales and similar fluidvelocities as observed at ODP Site 1130) to investigate thepotential error introduced by the 1-D assumptions. Tocompare the results of the two models, we plot a 1-D crosssection of the 2-D model, against the results of a 1-D modelusing the same parameters as the 2-D model. Fig. 7 showthat the main differences between the two approaches are2found in the resulting concentration of SO 4and H 2 S.(Fig. 7). This implies that the 1-D model overestimatesthe volumetric sulfate reduction rate compared to the 2-Dmodel. However, the isotopic composition remains thesame. We therefore conclude that the total fluxes computedby our model may contain a certain error, but that the ratioof the fluxes should be accurate.The d 18 O SO4 2 data from ODP Site 1130 show a distinctiveincrease up to a maximum of 28.6‰ at 27.5 mbsf, andsubsequently decrease to 11.6‰ at 300 mbsf (Fig. 8). Thereturn to values close to the seawater sulfate isotope compositionis mostly a function of the upwelling brine supplyingsulfate with a d 18 O SO4 2 11&. We will first explorewhether this d 18 O SO4 2 signal can be explained with a kineticisotope effect alone. The suggested ratio between thefractionation factors for O and S vary from 1:1.4 to 1:4

4228 U.G. Wortmann et al. / Geochimica et Cosmochimica Acta 71 (2007) 4221–4232mbsf [m]0255075100125150175200225250275300-50 0 50 100SO 4 [mM] 2-DH 2 S [mM] 2-Dδ 34 S SO 4 [ 0 / 00 ] 2-Dδ 34 S H 2 S [ 0 / 00 ] 2-DΔ 34 S [ 0 / 00 ] 2-DSO 4 [mM] 1-DH 2 S [mM] 1-Dδ 34 S SO 4 [ 0 / 00 ] 1-Dδ 34 S H 2 S [ 0 / 00 ] 1-DΔ 34 S [ 0 / 00 ] 1-DFig. 7. Comparison between the results of a hypothetical 2-Dmodel with a strong horizontal flow component (lines), and theresults of 1-D model which only considers the vertical flowcomponents of the 2-D model (symbols). The largest differences areseen for the concentrations of hydrogen sulfite and sulfate,suggesting that a 1-D model overestimates biological activity inthe presence of a lateral flow component. Note however that thedifferences are small, and that the difference in the isotope data isnegligible. The above results were obtained from models which usesimilar horizontal and vertical scales as observed in the westerntransect of ODP Leg 182 (i.e., Sites 1130 and 1132 Feary et al.,2000a). The models used a fractionation factor of a = 1.07, the 2-Dmodel used horizontal flow velocity of x x = 6.867 · 10 11 m/s, andboth models use a vertical flow velocity of x z = 3.173 · 10 11 m/s,and a sedimentation rate of 7.93 · 10 12 m/s.(Mizutani and Rafter, 1969; Aharon and Fu, 2003). Thefractionation factors for S reported for ODP Site 1130range from 75‰ (direct isotopic difference of dissolved2SO 4and H 2 S) to 65‰ (determined from reaction transportmodeling, Wortmann et al., 2001). We thus use a minimumfractionation factor of 16.25‰, and a maximumfractionation factor of 57‰ to investigate the theoreticald 18 O SO4 2 evolution with depth if it were caused by a kineticisotope effect associated with the volumetric sulfate reductionrate f. Note that the latter fractionation factor ford 18 O SO4 2 is substantially higher than any reported in the literature(up to 21‰, see Böttcher et al., 2005), and we donot imply that those fractionation factors do exist. Wemerely use this number to explore the upper solution envelopeof the model.The volumetric SRR is determined from inverse modelingof the sulfate concentration profile (see, Wortmann,2006). The rates shown in Fig. 8 are higher than those presentedby Wortmann (2006) as the model considers now amajor change in the sedimentation regime during the last20 ky (Hine et al., 2002) which affects the way the upwellingvelocity of the brine is calculated. The calculated SRR isthen used to model the kinetic isotope effect on d 18 O SO4 2 .Fig. 8 shows that a fractionation factor of 16.25‰ is toosmall to explain the observed d 18 O SO4 2 variations. Whilehigher fractionation factors are able to reproduce the peakvalue of d 18 O SO4 2 ¼ 28:6& at 27.5 mbsf, they fail to explainthe values observed in the middle of the profile.Increasing the fractionation factor to values greater than57‰ would solve the discrepancy between 250 and 60 mbsf,but must result in much higher d 18 O SO4 2 values between 60and 0 mbsf (see Fig. 8). An increase of the SRR in the upper30 mbsf, a possibility we cannot exclude, would also increasethe peak for the 16‰ model, but this scenario wouldstill be unable to explain the d 18 O SO4 2 data at 160 mbfs.The only way to explain the observed d 18 O SO4 2 values witha kinetic model would be a variable fractionation factor.However, the required kinetic fractionation factors aremuch higher than those reported in the literature (up to21‰, see Böttcher et al., 2005). We thus conclude that kineticfractionation of oxygen isotopes during microbial sulfatereduction is an unlikely explanation for the ODP Site1130 data.The d 18 O SO4 2 values may be also be influenced by oxygenisotope exchange reactions. Under typical marine conditions,oxygen isotope exchange rates are slow (10 7 years,Chiba and Sakai, 1985). However, metabolic intermediateslike APS or sulfite facilitate oxygen exchange reactions insidethe cytoplasm (Fritz et al., 1989). In this scenario, theoverall oxygen isotope effect depends on the volumetric exchangeflux b, on the completeness of the exchange reaction(k), and the isotopic equilibration fractionation factor .While we have no means to determine k from field measurements,data from ODP Site 1130 clearly shows that at30 mbsf, the isotopic signature of the interstitial water iswithin error similar to the empirical equilibrium constantas determined by Fritz et al. (1989). This implies that thecombination of k and b is large enough to achieve a completeexchange at this depth. We therefore assume thatk = 1, and acknowledge that we may underestimate b.Our model was built without a priory assumptionswhich of the metabolic intermediates facilitates the oxygenisotope exchange, and we are therefore unable to differentiatebetween the individual contributions of APS and sulfite.The measured d 18 O SO4 2 data reaches values of up to 28.6‰,whereas in experiments with sulfur disproportionating bacteriathe oxygen isotope exchange between water and sulfiteand subsequent reoxidation to sulfate typically results inmuch lower values off up to 21‰ at 28 to 35 °C (Böttcheret al., 2001, 2005). Assuming that the oxidation of sulfiteto sulfate adds one oxygen from the cytoplasmic water(d 18 O=0‰), the initial oxygen isotope equilibrium betweensulfite and water in the above experiments wouldhave been close 28.6‰.To achieve the observed d 18 O SO4 2 ¼ 28:6& through theaddition of a water-derived oxygen requires that either theinitial sulfite–water equilibrium was 38.1‰ or that theadded oxygen was enriched relative to the ambient water.In the case of sulfate reducing bacteria, the most likelysource of the oxygen used for the oxidation of sulfite toAPS is adenosine monophosphate (AMP, see Fig. 9).AMP is used in the transformation of sulfite to APS, a reactionthat is reversibly catalyzed by the flavoenzyme APSreductase (Fritz et al., 2002). While we do not know thed 18 O of AMP, its oxygen isotope composition is likely similarto the oxygen isotope equilibrium isotope fractionationbetween inorganic phosphate and water. Following

Oxygen isotope biogeochemistry of pore water sulfate... 4229Measured dataδ 18 O SO4 [ 0 / 00 VSMOV]δ 18 O IW [ 0 / 00 VSMOW]Modeled dataTheoretical δ 18 O SO4 equilibrium [ 0 / 00 VSMOV]δ 18 O SO4 exchange [ 0 / 00 VSMOV]δ 18 O SO4 kinetic α=1.01625 [ 0 / 00 VSMOV]δ 18 O SO4 kinetic α=1.057 [ 0 / 00 VSMOV]Net Flux SO 42Gross Flux SO 4 2- ε=29 0 / 00Gross Flux SO 4 2- ε=50 0 / 00020406080100mbsf1201401601802002202402602803000 10 20 30 40δ 18 O [ 0 / 00 VSMOV]10 -12 10 -11 10 -10mol/m 3 s -1Fig. 8. Oxygen isotope composition of the interstitial water (d 18 O IW ), the d 18 O values from dissolved sulfate (d 18 O SO4 ), as well as the results ofdifferent model runs assuming either a kinetic fractionation ðd 18 O SO4kinetic Þ process or fractionation by isotope exchange reactions d 18 O SO4exchange .The theoretical d 18 O SO4 2 equilibrium values were calculated using the equation of Fritz et al. (1989) from the shipboard temperature data andthe porewater d 18 O measurements.Coleman et al. (2005) and references therein, the oxygenisotopic equilibrium between water and phosphate at 5 °Ccan be calculated as (111.4 5)/4.3 = 24.7. This would implythat the initial sulfite–water equilibrium was 29.9‰,close to the equilibrium constant derived above from thedata reported by Böttcher et al. (2001, 2005).The net flux (i.e., the sulfate reduction rate, SRR) andthe gross flux (i.e., the exchange flux see Fig. 8) is on average14 times higher than the SRR. Similarly high rates arealso reported by Turchyn et al. (2006). These flux rates arehowever critically dependent on the choice of . We thereforealso modeled the case with a considerably higher equilibriumfractionation factor ( =50‰). Such an apparentconstant could result from a combination of an oxygen isotopeexchange effect with a additional kinetic oxygen isotopefractionation (hypothesis C). In this case, therequired volumetric exchange flux b is reduced to the pointthat it becomes smaller than the SRR in the upper 80 mbsf,and is only 4 times higher than the SRR further downcore.However, in the absence of data supporting such highapparent equilibrium constants, we consider a simple exchangemodel the better explanation.The fact that the calculated exchange flux is much greaterthan the flux produced by dissimilatory sulfate reductionposes, however, some difficulties for our understanding of2how SO 4is transported through the cell membrane.While Cypionka (1989) convincingly showed that large2and fast back-fluxes of SO 4across the cell membraneare possible, Brüchert (2004) argues that a large back-fluxwould reduce the membrane potential dW as the sulfateion carries a charge. We note however, that dW dependsnot only on the back-fluxes alone, but on the difference betweenthe sum of the forward and backward fluxes. FurthermoreCypionka (1989) showed that sulfate transportacross the cell membrane can be electroneutral if cell-externalsulfate concentrations are high enough.

4230 U.G. Wortmann et al. / Geochimica et Cosmochimica Acta 71 (2007) 4221–4232Fig. 9. The most likely source of the oxygen used for the oxidation of sulfite to APS is adenosine monophosphate (AMP). AMP is used in thetransformation of sulfite to APS, a reaction that is reversibly catalyzed by the flavoenzyme APS reductase (Fritz et al., 2002).mbsf050100150200250300350400FeP [wt%]0.000 0.025 0.050 0.075 0.1000.0 0.1 0.2 0.3 0.4FeT [wt\%]Fig. 10. Iron and pyrite concentration data for ODP-Site 1130.FeT, total iron content; FeP, pyrite bound iron.From a modeling point of view, we cannot discriminatebetween exchange processes related to sulfate reduction(hypothesis B and C), sulfur disproportionation (D), or directsulfide oxidation to sulfate (E). The latter two processes,however, depend on the availability of an oxidantto oxidize H 2 S to elemental sulfur. While it is currently unclearwhich substance could act as an oxidant at hypersulfidicSite 1130 (Wortmann et al., 2001), we can calculate themolar equivalent needed to sustain the above exchangefluxes, using Eq. (16). Assuming a 2:1 ratio between oxidantand oxidized sulfide would require 3000 mol/m 3 of oxidantat 30 mbsf. If the oxidizing phase were Fe 2 O 3 , thiswould be equivalent to an iron content of 14 wt%. If we allowfor a higher value, the required exchange fluxes wouldbe much smaller, e.g., with an =50‰, the exchange fluxwould be 4-fold, equivalent to 4 wt% Fe 2 O 3 . The iron concentrationdata shows that ODP Site 1130 contains notmore than 0.4 wt% Fe, which is about one order of magnitudeless than the amount discussed above (see Fig. 10).3.1. ConclusionsWe present d 18 O SO4 2 data from interstitial watersamples of ODP Site 1130. The maximum observedd 18 O SO4 2 values coincide within error with the experimentallydetermined steady state equilibrium value of 29‰ at5 °C. Reaction transport modeling shows that it is difficultto explain the d 18 O SO4 2 of ODP Site 1130 invoking kineticfractionation effects. If we consider however isotopic exchangereactions between metabolic intermediates andambient water, the measured data can be explained in aconsistent way with an oxygen isotope equilibrium fractionationfactor =29‰. Our model was built without a prioryassumptions which of the metabolic intermediates facilitatesthe oxygen isotope exchange, and we are therefore unableto differentiate between the individual contributions ofAPS, sulfite, and AMP. However, in a sulfite exchange scenariowith subsequent oxidation to sulfate with oxygen derivedfrom water, the oxygen isotope equilibrium betweensulfite and water would be at least 38‰, much higher thanany reported estimates so far. Therefore, we suggest thatthe oxygen isotope exchange either occurs between APS

Oxygen isotope biogeochemistry of pore water sulfate... 4231and water, or that sulfite is oxidized to sulfate with enrichedoxygen derived from AMP. The latter solution requires thatthe oxygen isotopic sulfite–water equilibrium value isaround 30‰, which is close to the value implied by otherstudies. We cannot fully discount the possibility that the exchangereaction occurs with a metabolic intermediate in theoxidative part of the sulfur cycle, but we show that thiswould require up to 3000 mol/m 3 of oxidant, which isinconsistent with the sediment geochemical data.The calculated volumetric oxygen isotope exchangefluxes are on average 14 times higher than the correspondingSRR. This suggests that only a small fraction of the sulfateentering the cell is used to gain metabolic energy. 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