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Geochimica et Cosmochimica Acta 71 (2007) 4221–4232

Oxygen isotope biogeochemistry of pore water sulfate in thedeep biosphere: Dominance of isotope exchange reactions with

ambient water during microbial sulfate reduction (ODP Site 1130)

Ulrich G. Wortmann a,*, Boris Chernyavsky a, Stefano M. Bernasconi b,Benjamin Brunner c, Michael E. Bottcher d, Peter K. Swart e

a Geobiology Isotope Laboratory, Department of Geology, University of Toronto, 22 Russellstr., Toronto, Ont., Canada M5S 3B1b ETH-Zurich, Geological Institute, 8092 Zurich, Switzerland

c Jet Propulsion Laboratory, California Institute of Technology, USAd Leibniz Institute for Baltic Sea Research, Seestr. 15, D-18119 Warnemunde, Germany

e Marine Geology and Geophysics, Rosenstiel School of Marine and Atmospheric Sciences Miami, FL, USA

Received 2 October 2006; accepted in revised form 12 June 2007; available online 27 June 2007

Abstract

Microbially mediated sulfate reduction affects the isotopic composition of dissolved and solid sulfur species in marine sed-iments. Experiments and field data show that the d18OSO4

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-inter-nal exchange reactions between enzymatically-activated sulfate (APS), and/or sulfite with cytoplasmic water. The isotopic fin-gerprint of 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 d18OSO4

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 d18OSO4

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 monophos-phate (AMP) during APS formation. Model calculations show that cell external reoxidation of reduced sulfur species wouldrequire up to 3000 mol/m3 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 d18OSO4

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.� 2007 Elsevier Ltd. All rights reserved.

1. INTRODUCTION

Microbial sulfate reduction is the major pathway of or-ganic matter (OM) oxidation in coastal-marine and conti-

0016-7037/$ - see front matter � 2007 Elsevier Ltd. All rights reserved.

doi:10.1016/j.gca.2007.06.033

* Corresponding author.E-mail address: uli.wortmann@utoronto.ca (U.G. Wortmann).

nental-shelf sediments (Jørgensen, 1982), and is afundamental process linking the geochemical cycles of car-bon, sulfur, and oxygen (e.g., Berner, 1982; Schidlowskiet al., 1983; Garrels and Lerman, 1984; Wortmann andChernyavsky, 2007). Sulfate-reducing microorganisms re-duce SO4

2� according to the following net reaction:

SO42� þ 2CH2O! H2Sþ 2HCO3

� ð1Þ

4222 U.G. Wortmann et al. / Geochimica et Cosmochimica Acta 71 (2007) 4221–4232

Although several details of the fractionation process re-main controversial, 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 back-ward reactions (Rees, 1973; Bruchert, 2004; Brunner andBernasconi, 2005). Culture experiments with dissimilatorysulfate reducers and field data show that the evolution ofthe d18OSO4

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; Bottcher et al., 1998, 1999; Brunner et al., 2005).Thus, if water is strongly depleted in 18O compared to sul-fate, 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). How-ever, 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 demon-strated that isotope exchange reactions between sul-fate and water are the dominant control factor ofthe d18OSO4

2� value (e.g., Mizutani and Rafter,1973; Fritz et al., 1989; Bottcher et al., 1998; Brunneret al., 2005; Knoller et al., 2006). However, the possi-bility of a kinetic d18OSO4

2� fractionation is still dis-cussed in the literature describing marineenvironments. This interpretation is based on theobservation that in certain studies the measuredd18OSO4

2� and d34S values of dissolved sulfate showa linear correlation (e.g. Aharon and Fu, 2000; Bott-rell et 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.

Cytoplasmic Membrane

Cytoplasm

SO42— SO42—

H2OH2O

APS

AMP

f1 f2 f3

b1 b2 b3

Fig. 1. The major fractionation steps and associated fluxes during microsum of the individual steps and the ratio between the forward and backw

(B) The d18OSO42� is a function of an oxygen isotope

exchange between metabolic intermediates and cyto-plasmic water during microbially-mediated sulfatereduction. Fritz et al. (1989) observed that the d34Sand d18OSO4

2� values of aqueous sulfate in a groundwater environment initially increased together, butthat the d18OSO4

2� value asymptotically approacheda constant value whereas the d34S value continuedto increase. This observation agrees with results fromanoxic pore waters of marine sediments (Zak et al.,1980; Bottcher et al., 1998, 1999, 2001). It was shownexperimentally that the final steady state d18OSO4

2�

value depends on the d18O value of the ambient water(Mizutani and Rafter, 1973; Fritz et al., 1989; Brun-ner et al., 2005) suggesting an isotope exchange reac-tion between ambient water and sulfate. Theexperimentally determined steady state value for bac-terial cultures (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 sug-gested that this exchange must take place betweenenzymatically activated sulfate (adenosine phospho-sulfate, APS) or sulfite and cytoplasmic water (Mizu-tani and Rafter, 1973; Fritz et al., 1989).

(C) The d18OSO42� 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 equilib-rium isotope value and create an apparent equilib-rium isotope factor, which would be greater thanthe actual equilibrium factor.

SO32— H2S H2S

f4 f5

b4 b5

bial reduction of sulfate. The overall isotope effect depends on theard fluxes. Modified after Brunner and Bernasconi (2005).

Oxygen isotope biogeochemistry of pore water sulfate. . . 4223

(D) d18OSO42� isotope effects are mainly a product of the

oxidative part of the sulfur cycle where H2S is oxi-dized to S0 and then disproportionated to SO4

2�

and sulfide. The net reaction describing this processcan be written as

4H2Oþ S0 ! 3H2Sþ SO42� þ 2Hþ ð2Þ

Bottcher et al. (2001, 2005) observed that the abovereaction results in a net isotope effect for d18OSO4

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 oxy-gen isotope exchange with ambient water. This processhas been used to explain the measurements from inter-stitial waters (e.g. Blake et al., 2006; Turchyn et al.,2006) and to explain the d18OSO4

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 d18OSO42� data for dissolved

SO42� from 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 background

The Great Australian Bight (GAB) forms the centralembayment of Australia’s southern continental margin, lo-cated between 124�E and 134�E and 32�S and 34�S. It rep-resents the 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 cool-

2000

1000

200

127 E

34 S

33 S

EyreJerboa

Terrace

100

GAB-13B(alternate)

1128

1130

1134

1132

Fig. 2. Locations of the sites drilled during ODP-Leg 182. Contour line(2000a).

water carbonate 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 seawa-ter (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 sig-nature of 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 SO4

2� from below, where SO42� concentrations increase

with depth until they stabilize at 300 mbsf at 65 mM (Fearyet al., 2000a). Modeling the conservative chloride ion con-centration data of Site 1130 shows that Site 1130 is influ-enced by strong upward advective flow on the order of2 mm/year (Wortmann, 2006). Inverse reaction-transportmodeling of the sulfate concentration data of this site, sug-gests that volumetric sulfate reduction rates (SRR) vary be-tween 600 and 65 pmol/cm3/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 ap-proaches 70‰ (see Fig. 4, and for a detailed discussion seeWortmann et al. (2001) and Wortmann (2006)).

4000

3000

500

129 E128 E

-1

0 50 km

1131

4500

1127

1126

1129

1133

s are in m and show water depth. Figure taken from Feary et al.

567 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.

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300

250

200

150

100

50

0 SO4 [mM] DataH2S [mM] Dataδ34S SO4 [0/00 VCDT] Dataδ34S H2S[0/00 VCDT] Data

-50 0 50

[mM]

-50 0 50

[ VCDT]

300

250

200

150

100

50

0

Dep

th [m

]

SO4 [mM] ModelH2S [mM] Modelδ34S SO4 [0/00 VCDT] Modelδ34S H2S [0/00 VCDT] Model

0/00

Fig. 4. Wortmann et al. (2001) reported d34S data from ODP Site 1130 which showed an isotopic difference of more than 70‰ between co-existing dissolved H2S and SO4

2�. 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 non-steady state parametrization used throughout in this paper, and that Fig. 7 demonstrates that the observed d34S-signatures are not an artifactof the 1-D modeling approach.

4224 U.G. Wortmann et al. / Geochimica et Cosmochimica Acta 71 (2007) 4221–4232

However, the interstitial water data of the upper 30 mbsfat ODP Site 182 are difficult to explain within the frame-work of a steady state diagenetic model. Previous publica-tions (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 Bight

sediments might be contourites. Thus we assume that thesediments at Site 1130 are redeposited older sediments,which provides a plausible mechanism for seawater irriga-tion, and allows us for the first time, to successfully modelthe upper 30 mbsf of OPD Site 1130 (see Fig. 5).

2. METHODS

The interstitial water samples were taken on board theJOIDES Resolution following the procedures given in

50

40

30

20

10

0

mbs

f

100 101 102 103 104 105

Age [yrs B.P.]

Site 1130 14C dataSite 1130 new age model

26 cm/ky

2.86 cm/y0 cm/y

500 1.000 1.500

50

40

30

20

10

0

mbs

f [m

]

New age modelOld age model

500

Cl- [mM]

50

40

30

20

10

0

Cl-

Site 1130 old age model

3.12 cm/y

0 cm/y

Fig. 5. Comparison between the Site 1130 14C 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. 14C age data from Hine et al. (2002).

Oxygen isotope biogeochemistry of pore water sulfate. . . 4225

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 H2S and inhi-bit further 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 asBaSO4 by the addition of BaCl2 within one hour after acid-ification. Samples were centrifuged, washed with hot water,and dried. For d18OSO4

2� measurements, about 135 lgBaSO4 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, sepa-rated on a Molsieve 5A, and subsequently measured on aThermo Finnigan Mat 253 in continuous flow mode usingthe Conflo III open split interface. The system was cali-brated by 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 repro-ducibility of the measurements was determined by runningseveral replicates of an in-house standard (BaSO4 withd18O = 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 H2O from interstitialwater samples were made in the Division of Marine Geol-ogy and Geophysics at the University of Miami using awater equilibration system (WEST) attached to an EuropaGEO. This setup determines the oxygen isotopic composi-

tion from CO2 which has been injected into serum bottles(�5 cm3) at slightly above atmospheric pressure contain-ing 0.5 cm3 of sample. The samples are subsequently equil-ibrated at 40 �C for 12 h without shaking. The process isentirely automated with the CO2 being injected and re-trieved using 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.

d18OSO42� ¼ 29:772� T� 0:1599 ð3Þ

where T denotes the shipboard measured temperature pro-file of OPD Site 1130 (Feary et al., 2000a), which weapproximated as

T ¼ 10:5þ z� 0:037 ð4Þ

where z denotes the depth in meters below seafloor.Total iron of selected samples was measured by X-ray

fluorescence 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 sul-phur (TRIS) which is considered to essentially representpyrite sulfur, was extracted from the sediments accordingto the one-step hot Cr(II)Cl2 distillation method (Fossingand Jørgensen, 1989) where the H2S was trapped quantita-tively as Ag2S in an AgNO3 solution and quantified

4226 U.G. Wortmann et al. / Geochimica et Cosmochimica Acta 71 (2007) 4221–4232

gravimetrically (Bottcher et al., 2004). Pyrite-iron contentswere calculated from TRIS contents using the ideal stoichi-ometry of pyrite. Iron, leachable with buffered Na-dithio-nite solution (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 formulation

The distribution of dissolved interstitial water species af-fected by diagenetic or biologically mediated conversionprocesses can be described as a process which involvestransport by diffusion, transport by advection, and a con-sumption or production function (Berner, 1964, 1980;Boudreau and Westrich, 1984; Boudreau, 1996). The stan-dard diagenetic equation (Berner, 1980) for a dissolved spe-cies describes the change of concentration with depth andtime as a function of diffusion, advection, porosity, andconsumption or production:

oðuCÞot

����z

¼o DB

oðuCÞoz þ uðDi þ DÞ oC

oz

h ioz

� oðuxCÞoz

� uf ð5Þ

where: C is the concentration, D are diffusion coefficients todescribe diffusion due to bioturbation (DB), diffusion due toirrigation (Di) , 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 pro-duction function is termed f.

In the following, we will assume that there is no biotur-bation, no diffusive irrigation, and that porosity changesdowncore only as a result of compaction (for a more de-tailed discussion see, Berner, 1980; Wortmann, 2006; Cher-nyavsky and Wortmann, 2007). Thus we can simplify Eq.(5) and write

uoCot¼ u

oCoz

oDozþ uD

o2C

oz2þ D

ouoz

oCoz� ux

oCoz� uf ð6Þ

These 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 SO4

2� is 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 Wort-mann et al. (2001), here we include two major sedimenta-tion events 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 sec-ond event, depositing 9.36 m of sediment within 300 yearsabout 200 years ago. These changes increase the estimatedupwelling velocity twofold to 1.2 · 10�10m/s, and allow usto successfully model the chemical profiles in the upper30 mbsf, which was not possible previously (see e.g., Wort-mann et al., 2001; Wortmann, 2006). All modeling wasdone with REMAP (Chernyavsky and Wortmann, 2007)which was modified to include a module to calculate isotopeexchange reactions.

2.2. Modeling the kinetic oxygen isotope effects in sulfate

In order to model kinetic oxygen isotope effects in sul-fate we assume that this effect is proportional to the volu-metric net reduction rate f. We thus express the isotopeeffect similar to the approach of Jørgensen (1979), and write

f ðS16O4Þ ¼a½S16O4�

½SO4� þ ða� 1Þ½S16O4�f ð7Þ

f ðS18O4Þ ¼½S18O4�

a½SO4� � ða� 1Þ½S18O4�f ð8Þ

where values in square brackets denote concentrations.Note, that S18O is merely a notation to describe the concen-tration of 18O in a sulfate molecule but in no way impliesthe existence of a molecule with 4 18O isotopes (whichwould be extremely rare under natural conditions). Com-bining Eqs. (5) and (7), we obtain

uo½S16O4�

ot¼u

o½S16O4�oz

oDozþuD

o2½S16O4�

oz2þD

ouoz

o½S16O4�oz

�uxo½S16O4�

oz�u

a½S16O4�½SO4�þ ða� 1Þ½S16O4�

f ð9Þ

and similarly for 18O. This allows us to describe the evolu-tion of the d18OSO4

2� values as a function of depth, a and f.

2.3. Modeling the oxygen isotope exchange between water

and sulfate

In 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 iso-tope fractionation during microbial sulfate reduction, i.e.,that the d18OSO4

2� value depends only on the exchange reac-tion between water and SO4

2�.As equilibration reactions require zero net flow, we can

describe 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 en-ters a 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 de-scribe the 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 d18OSO4

2� value, the flux of 16O and 18O isproportional to the initial ratio of the given isotope to thetotal sulfate amount

f ðS16O4Þ ¼ f½S16O4�½SO4�

ð10Þ

and similarly for f (18O).We assume that there is no isotope effect on the

d18OSO42� composition for the flux associated with the ex-

change reaction and we can write this flux as being propor-tional to the initial ratio of the given isotope to the totalsulfate amount as in Eq. (10). Since the total concentration

Sulfate inporewater

sulfate reducing bacteria

sulfate reduction {f } • δ18OSO4 porewater

exchange flux {b} • δ18OSO4 porewater

δ18OSO4 cytoplasm {δOut} =

δ18OSO4 porewater {δIn} • (1-k) + k • (δH2O + ε)

exchange flux {b} • δ18OSO4 cytoplasm

Fig. 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.

Oxygen isotope biogeochemistry of pore water sulfate. . . 4227

of SO42� is not changed by the exchange reaction, we have

the additional constraint that the sum of the fluxes of 18Oand 16O cannot add any new sulfate, i.e., they must cancel

bðS16O4Þ þ bðS18O4Þ � 0 ð11Þ

where 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 as

bðS16O4Þ ¼ b � ½S16O4�½SO4�

þ 1

1þ ðdout=1000þ 1ÞR0

� �ð12Þ

bðS18O4Þ ¼ b � ½S18O4�½SO4�

þ ðdout=1000þ 1ÞR0

1þ ðdout=1000þ 1ÞR0

� �ð13Þ

where R0 refers to the isotopic ratio of V-SMOW, and

dout ¼ dinð1� kÞ þ kðdH2O þ �Þ ð14Þ

and dH2O denotes the oxygen isotopic composition of thepore-water H2O. The reaction transport model to describethe oxygen isotope exchange for 16O is then a combinationof Eqs. (6, 10 and 11)

uo½S16O4�

ot¼u

o½S16O4�oz

oDozþuD

o2½S16O4�oz2

þDouoz

o½S16O4�oz

�uxo½S16O4�

oz�uf ðS16O4ÞþubðS16O4Þ ð15Þ

and similarly for [S18O4].

2.4. Modeling cell external oxidation

Cell external oxidation of H2S 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 write

oðð1� uÞ½X s�Þot

¼ � oðð1� uÞxs½X s�Þoz

� ð1� uÞf� n ð16Þ

where Xs denotes a solid oxidant, f the gross flux of SO42�,

xs the sedimentation rate, and n the molar ratio of the oxi-dation reaction.

3. RESULTS AND DISCUSSION

Under 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 suf-ficiently constrained 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 arefound in the resulting concentration of SO4

2� and H2S.(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 d18OSO42� data from ODP Site 1130 show a distinc-

tive increase 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 com-position is mostly a function of the upwelling brine supply-ing sulfate with a d18OSO4

2� � 11&. We will first explorewhether this d18OSO4

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

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0m

bsf [

m]

SO4 [mM] 2-DH2S [mM] 2-Dδ34S SO4 [0/00] 2-Dδ34S H2S [0/00] 2-DΔ34S [0/00] 2-DSO4 [mM] 1-DH2S [mM] 1-Dδ34S SO4 [0/00] 1-Dδ34S H2S [0/00] 1-DΔ34S [0/00] 1-D

Fig. 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 xx = 6.867 · 10�11 m/s, andboth models use a vertical flow velocity of xz = 3.173 · 10�11 m/s,and a sedimentation rate of 7.93 · 10�12 m/s.

4228 U.G. Wortmann et al. / Geochimica et Cosmochimica Acta 71 (2007) 4221–4232

(Mizutani and Rafter, 1969; Aharon and Fu, 2003). Thefractionation factors for S reported for ODP Site 1130range from 75‰ (direct isotopic difference of dissolvedSO4

2� and H2S) to 65‰ (determined from reaction trans-port modeling, Wortmann et al., 2001). We thus use a min-imum fractionation factor of 16.25‰, and a maximumfractionation factor of 57‰ to investigate the theoreticald18OSO4

2� evolution with depth if it were caused by a kineticisotope effect associated with the volumetric sulfate reduc-tion rate f. Note that the latter fractionation factor ford18OSO4

2� is substantially higher than any reported in the lit-erature (up to 21‰, see Bottcher et al., 2005), and we donot imply that those fractionation factors do exist. Wemerely use this number to explore the upper solution enve-lope of the model.

The volumetric SRR is determined from inverse model-ing of the sulfate concentration profile (see, Wortmann,2006). The rates shown in Fig. 8 are higher than those pre-sented by 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 d18OSO4

2� .Fig. 8 shows that a fractionation factor of 16.25‰ is toosmall to explain the observed d18OSO4

2� variations. Whilehigher fractionation factors are able to reproduce the peakvalue of d18OSO4

2� ¼ 28:6& at 27.5 mbsf, they fail to ex-

plain the 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 d18OSO4

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 in-crease the peak for the 16‰ model, but this scenario wouldstill be unable to explain the d18OSO4

2� data at 160 mbfs.The only way to explain the observed d18OSO4

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 Bottcher et al., 2005). We thus conclude that ki-netic fractionation of oxygen isotopes during microbial sul-fate reduction is an unlikely explanation for the ODP Site1130 data.

The d18OSO42� values may be also be influenced by oxygen

isotope exchange reactions. Under typical marine condi-tions, oxygen isotope exchange rates are slow (107 years,Chiba and Sakai, 1985). However, metabolic intermediateslike APS or sulfite facilitate oxygen exchange reactions in-side the cytoplasm (Fritz et al., 1989). In this scenario, theoverall oxygen isotope effect depends on the volumetric ex-change flux 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 measure-ments, 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 com-plete exchange 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 differenti-ate between the individual contributions of APS and sulfite.The measured d18OSO4

2� data reaches values of up to 28.6‰,whereas in experiments with sulfur disproportionating bac-teria the 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 (Bottcheret al., 2001, 2005). Assuming that the oxidation of sulfiteto sulfate adds one oxygen from the cytoplasmic water(d18O = 0‰), the initial oxygen isotope equilibrium be-tween sulfite and water in the above experiments wouldhave been close 28.6‰.

To achieve the observed d18OSO42� ¼ 28:6& through the

addition 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 reac-tion that is reversibly catalyzed by the flavoenzyme APSreductase (Fritz et al., 2002). While we do not know thed18O of AMP, its oxygen isotope composition is likely sim-ilar to the oxygen isotope equilibrium isotope fractionationbetween inorganic phosphate and water. Following

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mol/m3 s-1

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Gross Flux SO42- ε=29 0/00

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Measured data

Modeled data

Fig. 8. Oxygen isotope composition of the interstitial water (d18OIW), the d18O values from dissolved sulfate (d18OSO4), as well as the results of

different model runs assuming either a kinetic fractionation ðd18OSO4kineticÞ process or fractionation by isotope exchange reactions d18OSO4exchange

.The theoretical d18OSO4

2� equilibrium values were calculated using the equation of Fritz et al. (1989) from the shipboard temperature data andthe porewater d18O measurements.

Oxygen isotope biogeochemistry of pore water sulfate. . . 4229

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 im-ply that the initial sulfite–water equilibrium was 29.9‰,close to the equilibrium constant derived above from thedata reported by Bottcher 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 aver-age 14 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 there-fore also modeled the case with a considerably higher equi-librium fractionation factor (� = 50‰). Such an apparentconstant could result from a combination of an oxygen iso-tope exchange effect with a additional kinetic oxygen iso-tope fractionation (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 ex-change model the better explanation.

The fact that the calculated exchange flux is much great-er than the flux produced by dissimilatory sulfate reductionposes, however, some difficulties for our understanding ofhow SO4

2� is transported through the cell membrane.While Cypionka (1989) convincingly showed that largeand fast back-fluxes of SO4

2� across the cell membraneare possible, Bruchert (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 be-tween the sum of the forward and backward fluxes. Fur-thermore Cypionka (1989) showed that sulfate transportacross the cell membrane can be electroneutral if cell-exter-nal sulfate concentrations are high enough.

Fig. 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).

4230 U.G. Wortmann et al. / Geochimica et Cosmochimica Acta 71 (2007) 4221–4232

From a modeling point of view, we cannot discriminatebetween exchange processes related to sulfate reduction(hypothesis B and C), sulfur disproportionation (D), or di-rect sulfide oxidation to sulfate (E). The latter two pro-cesses, however, depend on the availability of an oxidantto oxidize H2S to elemental sulfur. While it is currently un-

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Fig. 10. Iron and pyrite concentration data for ODP-Site 1130.FeT, total iron content; FeP, pyrite bound iron.

clear which substance could act as an oxidant at hypersulfi-dic Site 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/m3 of oxi-dant at 30 mbsf. If the oxidizing phase were Fe2O3, thiswould be equivalent to an iron content of 14 wt%. If we al-low for 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% Fe2O3. The iron con-centration data shows that ODP Site 1130 contains notmore than 0.4 wt% Fe, which is about one order of magni-tude less than the amount discussed above (see Fig. 10).

3.1. Conclusions

We present d18OSO42� data from interstitial water

samples of ODP Site 1130. The maximum observedd18OSO4

2� values coincide within error with the experimen-tally determined steady state equilibrium value of 29‰ at5 �C. Reaction transport modeling shows that it is difficultto explain the d18OSO4

2� of ODP Site 1130 invoking kineticfractionation effects. If we consider however isotopic ex-change reactions between metabolic intermediates andambient water, the measured data can be explained in aconsistent way with an oxygen isotope equilibrium fraction-ation factor � = 29‰. Our model was built without a prioryassumptions which of the metabolic intermediates facili-tates the oxygen isotope exchange, and we are therefore un-able to differentiate between the individual contributions ofAPS, sulfite, and AMP. However, in a sulfite exchange sce-nario with subsequent oxidation to sulfate with oxygen de-rived from 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. . . 4231

and 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 ex-change reaction occurs with a metabolic intermediate in theoxidative part of the sulfur cycle, but we show that thiswould require up to 3000 mol/m3 of oxidant, which isinconsistent with the sediment geochemical data.

The calculated volumetric oxygen isotope exchangefluxes are on average 14 times higher than the correspond-ing SRR. This suggests that only a small fraction of the sul-fate entering the cell is used to gain metabolic energy. Theratio between the modeled fluxes appears to be depthinvariant over large intervals, and we hypothesize that thisbehavior is caused by decreasing cell densities with depth,while the cell-specific SRRs remain constant.

ACKNOWLEDGMENTS

Discussions with Laura Lee and James Walker, and thecomments of three anonymous reviewers helped to focus ourideas. M.E.B. thanks S. Lilienthal and the Max Planck Societyfor support. U.G.W. thanks the Natural Sciences and Engineer-ing Research Council Canada NSERC, which supported thisstudy, and the German Science Foundation DFG which sup-ported his participation on ODP Leg 182. U.G.W. and P.K.S.thank crew, scientific party, and technicians of the JOIDES

Resolution on ODP Leg 182 for their support and commitmentwhich facilitated the recovery of these samples under difficultconditions. The authors thank B. Schnetger (ICBM Oldenburg)for XRF analyses and M.E.B. thanks A. Schipper for technicalassistance.

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