Microbial sulfate reduction rates and sulfur and oxygen isotope fractionations at oil and gas seeps in deepwater Gulf of Mexico

Geochimica et Cosmochimica Acta, Vol. 64, No. 2, pp. 233–246, 2000 Copyright © 2000 Elsevier Science Ltd Printed in the USA. All rights reserved 0016-7037/00 $20.00 ⫹ .00

Pergamon

PII S0016-7037(99)00292-6

Microbial sulfate reduction rates and sulfur and oxygen isotope fractionations at oil and gas seeps in deepwater Gulf of Mexico PAUL AHARON* and BAOSHUN FU Department of Geology and Geophysics, Louisiana State University, Baton Rouge, Louisiana 70803, USA (Received March 22, 1999; accepted in revised form July 5, 1999)

Abstract—Sulfate reduction and anaerobic methane oxidation are the dominant microbial processes occurring in hydrate-bearing sediments at bathyal depths in the Gulf of Mexico where crude oil and methane are advecting through fault conduits to the seafloor. The oil and gas seeps are typically overlain by chemosynthetic communities consisting of thiotrophic bacterial mats (Beggiatoa spp.) and methanotrophic mussels (Bathymodiolus spp.), respectively. Cores were recovered with a manned submersible from fine-grained sediments containing dispersed gas hydrates at the threshold of stability. Estimated sulfate reduction rates are variable but generally are substantially higher in crude oil seeps (up to 50 times) and methane seeps (up to 600 times) relative to a non-seep reference sediment (0.0043 ␮mol SO2⫺ cm⫺3 day⫺1). Sulfur and oxygen isotope 4 fractionation factors are highest in the reference sediment (␣S ⫽ 1.027; ␣O ⫽ 1.015) but substantially lower in the seep sediments (␣S ⫽ 1.018 to 1.009; ␣O ⫽ 1.006 to 1.002) and are controlled primarily by kinetic factors related to sulfate reduction rates. Kinetic effects also control the ␦34S/␦18O ratios such that slow microbial rates yield low ratios whereas faster rates yield progressively higher ratios. The seep data contradict previous claims that ␦34S/␦18O ratios are diagnostic of either microbial sulfate reduction at a fixed ␦34S/␦18O ratio of 4/1 or lower ratios caused by SO4–H2O equilibration at ambient temperatures. The new results offer a better understanding of methane removal via anaerobic oxidation in the sulfate reduction zone of hydratebearing sediments and have significant implications regarding the origin and geochemical history of sedimentary sulfate reconstructed on the basis of ␦34S and ␦18O compositions. Copyright © 2000 Elsevier Science Ltd A known feature of sulfate reduction is the measured decline of dissolved SO4 coupled with the increase of dissolved H2S with depth in interstitial fluids. Another consequence of sulfate reduction is the enrichment of both 34S and 18O in the dissolved sulfate because bacteria discriminate to various degrees against 34 S and 18O isotopes (Harrison and Thode, 1958; Mizutani and Rafter, 1969). Studies reporting sulfur isotope fractionations caused by microbial sulfate reduction in natural settings are rare, and determinations of both oxygen and sulfur isotopes of the residual sulfate are even fewer. Exceptional are the studies of Zak et al. (1980) and Fritz et al. (1989) which compared sulfur and oxygen isotope fractionations in deep-sea sediments and landfill sites, respectively. Knowledge of isotope fractionations accompanying microbial sulfate reduction are important because both 34S and 18O concentrations have been used as indicators of the origin and geochemical history of the sulfate (Fritz et al., 1989). However, laboratory experiments using pure cultures failed to resolve the question whether or not the oxygen isotope fractionation is controlled either by kinetic or thermodynamic processes because the experiments were difficult to control and the reactions were claimed to be inhibited by bacterial poisoning with excess H2S (Lloyd, 1968; Mizutani and Rafter, 1969, 1973; Fritz et al., 1989). It is therefore of interest to assess the isotope fractionations in a relatively simple natural marine environment where microbial sulfate reduction and concomitant organic matter decomposition are dominant. Moreover, it is of importance to contrast microbial sulfate reduction rates and isotope fractionations in deepwater sediments harboring chemosynthetic organisms vis-a-vis shallow water photosynthetic sediments. In this study we explore the factors controlling isotope

1. INTRODUCTION

Dissolved sulfate is abundant in seawater (⬃28 mM/L) and its microbial reduction is one of the dominant processes in the early diagenesis of marine sediments (Goldhaber and Kaplan, 1974). Modern marine sedimentary environments sustaining intense microbial sulfate reduction include stratified inland seas (e.g., Black Sea, Deuser, 1970; Sweeney and Kaplan, 1980), fjords (e.g., Saanich Inlet, Nissenbaum et al., 1972), continental shelves (Jørgensen, 1982), organic-rich deltas (Lin and Morse, 1991), and hydrothermal sediments (Jørgensen et al., 1992). In all these cases organic carbon of either terrestrial and/or marine origin serves as the primary electron-donor and metabolic substrate for sulfate anaerobic respiration: SO42⫺ ⫹ 2(CH2O) ⫽ H2S ⫹ 2HCO3⫺.

(1)

Hydrogen sulfide derived from microbial sulfate reduction has also been recently documented in cold submarine seeps along active (e.g., Sagami Bay, Japan, Masuzawa et al., 1992) and passive (e.g., North Sea, Dando et al., 1991) margins where light aliphatic hydrocarbons, primarily methane, serve as the reduced carbon source (Aharon, 1999): SO42⫺ ⫹ CH4 ⫽ H2S ⫹ CO32⫺ ⫹ H2O.

(2)

In turn, the H2S produced within the sediment by free-living microbial communities serves as nourishment for chemosynthetic symbions living within the tissues of benthic fauna inhabiting the interface between underlying anoxic sediments and the overlying oxygen-rich bottom waters. * Author to whom correspondence should be addressed. 233

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Fig. 1. Bathymetric map of the northern Gulf of Mexico showing distribution of extant hydrocarbon seeps. The two sites explored in this study are indicated by arrows.

fractionations in seep sediments at bathyal depths and cold ambient temperatures in the Gulf of Mexico. For this purpose, sediment cores from hydrocarbon seeps on the Gulf of Mexico continental slope (Fig. 1) were acquired using a manned submersible, and chemical and isotope compositions of pore fluids were analyzed. The northern Gulf of Mexico (Fig. 1) is a seemly site to investigate the isotope effects of microbial sulfate reduction in seeps on account of (i) pervasiveness of hydrocarbon seepage at cold ambient temperatures (Roberts and Aharon, 1994); (ii) diversity of hydrocarbons advecting on the seafloor and ranging from crude oil, methane in gas and solid hydrate forms derived from biogenic and/or thermogenic sources (Aharon, 1994a), and (iii) highly diverse and abundant chemosynthetic biota (Aharon, 1994b). The objectives of this study are to (1) determine the sulfate reduction rates in a hydrocarbon-driven anoxic system; (2) derive the sulfur and oxygen isotope fractionations during microbial sulfate reduction; (3) explore the factors controlling the reduction rates and the extent of isotope fractionations in seeps, and (iv) provide additional insight on the poorly known mechanism of isotope fractionation during microbial sulfate reduction in natural settings by coupling of sulfur and SO4oxygen isotope systematics. 2. METHODS Push cores up to 50 cm long were taken in seep sediments harboring chemosynthetic communities on the northern Gulf of Mexico continental slope during 1995 submersible dives with Johnson-Sea-Link. After recovery, the cored sediments were brought to the laboratory on the support surface ship where continuous pore fluid profiles were acquired with a squeezer modified from Jahnke (1988). The pore fluids were passed into plastic disposable syringes inserted into the core barrel at 3– 4 cm depth intervals and were filtered through a 0.25 ␮m filter capping the syringe. Pore fluids were analyzed immediately after recovery for dissolved sulfide and sulfate by the colorimetric method of Cline (1969) and the barium gravimetric method (Presley, 1969), respectively. The analytical error (⫾1␴) based on replicate analyses of standards was 4.5% for sulfide and 0.5% for sulfate determinations. Barium sulfate precipitated from the pore fluids were used for isotope analyses. ␦34S of SO4 was measured by graphite reduction of BaSO4 to BaS, which was subsequently dissolved in water and precipitated as Ag2S with 0.1 N AgNO3. Ag2S was oxidized with Cu2O

(mixture 1:4 by mass) at 900°C to SO2, which was analyzed with the mass spectrometer. The overall error (⫾1␴) of ␦34S determinations is better than 0.2‰ and the samples are reported relative to the Canyon Diablo Triolite (CDT) reference standard. ␦18O of SO4 were analyzed in the manner described by Rafter and Mizutani (1967) and Mizutani (1971). Intimate mixtures of BaSO4 and spectrographically pure (99.999%) graphite (1:2 by mass) were decomposed in resistanceheated platinum boats near 1000°C. In addition to CO2 some CO was produced, which was converted into CO2 using a high-voltage electrical discharge between platinum electrodes in a trap assembly immersed in liquid nitrogen. Isotope measurements were made on a Nier-type triple collector gas source mass spectrometer. Accuracy and precision of ␦18O determinations is better than 0.3‰ based on repeat standard analyses and the results are reported relative to standard mean ocean water (SMOW). The ␦34S and ␦18O values of the OGS standard (seawater-derived sulfate precipitated as BaSO4) measured in this study are 20.3‰ and 9.62 ⫾ 0.09‰ (n ⫽ 5), respectively. Sulfate and sulfide concentrations and ␦34S and ␦18O compositions of pore fluids SO4 acquired from 10 push cores are listed in Table 1. Block name and numbers are the ones used by the Mineral Management Service (Fig. 1). Core sets 2635 (n ⫽ 2) and 2639 (n ⫽ 3) were taken in crude oil seeps overlain by orange–red colored bacterial mats (Beggiatoa spp.) at a water depth of 630 m in Green Canyon #232 (27°44⬘N; 91°18⬘W). Core set 2647 (n ⫽ 3) was recovered from methane seeps overlain by mussel beds (Bathymodiolus spp.) in Green Canyon #184/185 (Bush Hill) at a water depth of 590 m (27°46⬘N; 91°30⬘W). Core 2647-4, representing non-seep, reference sediment, was sampled about 10 m away from seeps in Green Canyon #184/185 at a water depth of 590 m. Except the reference core, all other cores were recovered from hydrate-bearing sediments (Roberts and Aharon, 1994) where gas hydrates (clathrates) crop out along crevasses. With the two notable differences below, seep sediments investigated here are dark-grey to black, fine-grained clay to silt size, which are typical of the siliciclastic-sediment apron draping the slope distal to the Mississippi River (Fig. 1). The differences from normal slope sediments are random stains of crude in oil seep sediments and sand to pebble size (up to 1 cm) carbonate concretions, constituting 40% to 60% of the sediment, commonly occurring in cores whose pore fluids show high levels of H2S (e.g., 2639-3 and 2639-4, Table 1). 3. RESULTS

3.1. Dissolved Sulfate and Sulfide Profiles of pore-fluid chemistry generally show an increase of sulfide concentrations downcore accompanied by a progressive decline in sulfate concentrations (Fig. 2A). Pore fluid

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235

Table 1. Chemical and isotope compositions of pore fluids from reference and hydrocarbon seep sediments. Sample Name

Core Depth (cm)

GOM Bottom Water 2647-4-1† 2647-4-2 2647-4-3 2647-4-4 2647-4-5 2647-4-6 2647-4-7 2647-4-8 2647-4-9 2647-4-10

⫺0.2 ⫺2.0 ⫺4.0 ⫺6.0 ⫺9.0 ⫺11.0 ⫺13.0 ⫺16.0 ⫺19.0 ⫺22.0

Salinity (‰) 38

pH* 7.7 7.6 7.9 7.9 7.5 7.8 7.7 7.7 7.7 7.5 7.5

2635-1-2 2635-1-3 2635-1-4 2635-1-5 2635-1-6 2635-2-1 2635-2-2 2635-2-3 2635-3-1 2635-3-2 2635-3-3

⫺5.0 ⫺8.5 ⫺12.5 ⫺16.0 ⫺19.0 ⫺4.0 ⫺7.0 ⫺10.0 ⫺0.5 ⫺3.5 ⫺6.5

38 38 38 38 38 38 38 38 38 38 38

7.6 7.6 7.6 7.8 7.9 7.4 7.4 7.5 7.4 7.4 7.4

2639-2-2‡ 2639-2-3 2639-2-4 2639-2-5 2639-2-6 2639-2-7 2639-2-8 2639-3-1 2639-3-2 2639-3-3 2639-3-4 2639-3-5 2639-3-6 2639-3-7 2639-3-8 2639-4-1 2639-4-2 2639-4-3 2639-4-4 2639-4-5 2639-4-6 2639-4-7 2639-4-8 2639-4-9

⫺2.5 ⫺5.0 ⫺8.0 ⫺11.0 ⫺14.5 ⫺18.0 ⫺21.0 ⫺3.0 ⫺5.0 ⫺7.0 ⫺10.0 ⫺12.0 ⫺14.0 ⫺17.0 ⫺20.5 ⫺3.0 ⫺5.0 ⫺8.0 ⫺10.0 ⫺13.0 ⫺15.0 ⫺18.0 ⫺21.0 ⫺23.0

37 38 38 38 38 38 38

7.8 8.0 8.2 8.1 8.2 8.3 8.3

§

⫺2.0 ⫺6.0 ⫺2.0 ⫺4.0 ⫺6.0 ⫺9.0 ⫺12.0 ⫺15.0 ⫺18.0 ⫺21.0 ⫺1.0 ⫺3.0 ⫺5.0 ⫺7.0 ⫺10.0 ⫺13.0 ⫺16.0 ⫺20.0 ⫺25.0

38 38 38 37 38 38 38 38 38 38

‡

2647-1-1 2647-1-2 2647-2-1 2647-2-2 2647-2-3 2647-2-4 2647-2-5 2647-2-6 2647-2-7 2647-2-8 2647-3-1 2647-3-2 2647-3-3 2647-3-4 2647-3-5 2647-3-6 2647-3-7 2647-3-8 2647-3-9

8.3 8.8 8.4 8.9 8.9 8.8 8.9 9.0 9.0 9.0

Cl (mM/L)

SO4 (mM/L)

(SO4/Cl) ⫻102

H2S (mM/L)

␦34S (SO4) (‰ CDT)

␦18O (SO4) (‰ SMOW)

568

29.0

5.1

20.3

9.7

20.7 21.1 21.3

10.5 11.0 11.1

23.2

12.0

25.4

0.0 0.0 0.0 0.0 0.0 0.7 0.8 0.7 0.8 0.9 1.0

23.4

12.1

22.8 22.0 21.1 18.8 15.9 26.3 25.7 24.1 28.6 27.3 26.9

3.9 4.2 5.1 7.0 8.6 1.7 2.1 3.3 0.2 1.0 1.3

23.1 24.9 26.7

12.6 13.0 13.9

26.2 34.8 35.7 30.4 39.8

13.3 17.4 16.7 14.8 17.8

30.7 35.6 36.6

16.1 17.0 17.4

26.9

14.5

31.6

15.3

39.0 44.2 36.8 44.2 49.4

18.9 20.5 18.6 20.1 22.4

28.2 27.7 27.5 25.8

582

558 573 580 578 585 581 583

575 567 576 578 575 570 577 573

19.4 17.7 15.2 14.8 10.8 8.7 7.8 22.8 12.3 14.3 18.1 11.8 1.2 1.0 0.3 22.0 19.5 16.4 15.0 13.2 23.3 20.1 18.4 16.4 11.3 4.4 14.3 5.3 2.6 0.4 0.5 0.3 0.4 1.0 16.8 15.5 12.2 8.0 5.0 1.9 1.7 1.2 1.0

* pH reported in Fu (1998). Core from hydrocarbon seep-free sediments in GC 184/185 at 590 m depth. ‡ Cores from oil seep sites in GC 232 at 624 m. § Cores from gas seep sites in GC 184/185 at 590 m depth. †

2.7

3.5 3.1 2.6 2.6 1.8 1.5 1.3

2.5 0.93 0.45 0.07 0.09 0.05 0.07 0.17

4.9 8.7 10.0 10.6 13.2 15.2 16.3 2.0 9.8 12.4 13.6 15.0 20.3 21.3 21.1 1.2 5.5 6.5 7.5 7.9 7.5 6.7 5.5 4.3 11.4 15.0 7.7 17.3 19.3 19.7 19.7 19.9 19.9 20.3 7.9 9.8 11.0 13.6 14.0 18.7 19.3 19.3 19.5

70.8 32.2 35.2 40.2 44.3 47.7 50.9

16.4 17.3 18.5 20.1 21.3 21.8

59.1

23.6

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Fig. 2. Sulfate and hydrogen sulfide concentrations in pore fluids of reference sediment and representative oil and gas seeps as a function of core depth (data from Table 1). (A) Dissolved sulfate (SO4); (B) hydrogen sulfide (H2S).

sulfate in the reference sediment ranges from 25.4 to 28.2 mM/L (Table 1) which is slightly lower than that of the ambient Gulf of Mexico bottom waters containing 28.9 mM/L dissolved sulfate. In contrast, pore fluids from oil and gas seeps are highly depleted in sulfate (down to 0.3 mM/L) relative to the bottom water. It is evident from the slopes of the sulfate profiles in Figure 2A that depletion rates in gas seeps are generally higher than those in oil seeps. The exception is core 2639-3 from an oil seep where the lower part of the core has similar sulfate concentrations as those measured in gas seeps at comparable depths (Table 1). Ambient bottom water contains sulfide below the detection limit of 1 ␮M/L. Pore fluids from the reference sediment contain sulfide at levels not exceeding 1 mM/L (Table 1). In contrast, H2S concentrations are highly elevated in all pore fluids from seeps (up to 20 and 21 mM/L for gas and oil seeps, respectively) and the levels typically increase downcore (Fig. 2B). Sulfide concentrations in gas seeps are generally higher than those in oil seeps at same core depths. The only exception is core 2639-3 from an oil seep where the lower parts of the core contain slightly higher sulfide levels than those in gas seeps at comparable depths (Table 1). The depletion of pore fluids SO4 accompanied by enrichment of H2S with depth (Fig. 2) indicates that sulfate reduction has occurred within the surficial seep sediments and the H2S was not produced through thermogenic processes in deep-seated reservoirs (Krouse, 1977). This observation is substantiated by the inverse linear relationship between sulfate and sulfide (Fig. 3) as expected from bacterial consumption of SO4 and concomitant production of H2S during anaerobic sulfate reduction (Eqns. 1 and 2 above). Since the consumption of one mole of SO4 by bacteria generates one mole of H2S during reduction, a line with a slope of 1 is expected in coordinates of SO4/H2S (Fig. 3). The observation that all pore fluid data fall below the line with the slope of 1 suggests that some process(es) removing H2S from pore fluids must occur. Likely processes are either consumption of H2S by thiotrophic consumers overlying the seeps and/or precipitation of sulfides within the sediment. Bacterial mats consist of a group of Beggiatoa spp. sulfide oxidizers which live at the seawater-sediment interface and obtain their energy

Fig. 3. Inverse linear relationship between SO4 and H2S in pore fluids from seeps. Seep data fall below the line with a slope of 1 suggesting partial removal of H2S from the pore fluids.

from oxidation of H2S yielding molecular sulfur (Larkin et al., 1994). Unlike the thiotrophic Bathymodiolus spp. mussels reported at hydrothermal vents (Rau and Hedges, 1979), the ones typical of hydrocarbon seeps in the Gulf of Mexico seem to contain primarily methanotrophic bacteria symbionts (Cary et al., 1988), although the co-occurrence of both thiotrophic and methanotrophic bacteria in the gills of the seep mussels has not been ruled out (Fisher et al., 1987). Sulfide precipitation may also occur because all seep sediments show dark gray to black color which could be caused by the presence of ferrous sulfide. Regardless whether the pathway of H2S removal is chemosynthetic, inorganic, or both, it can be estimated from Figure 3 that about 31% of the H2S produced in the seep sediment is removed from the pore fluids. 3.2. Sulfur and Oxygen Isotope Compositions Generally, both ␦18O and ␦34S values increase with depth in all the cores. The isotope compositions of sulfate in pore fluids from the reference sediment range from 10.5‰ to 12.1‰ (SMOW) and from 20.7‰ to 23.4‰ (CDT), respectively (Table 1). These values are slightly heavier than those of modern seawater (␦34S ⫽ 20.3‰ and ␦18O ⫽ 9.7‰, Faure, 1986). In sharp contrast, pore fluids from seep sediments are anomalously enriched in both 18O and 34S relative to modern seawater with ␦18O and ␦34S values reaching up to 23.6‰ and 70.8‰, respectively. Inverse relationships observed between the isotopes and SO4 concentrations (Table 1) are attributed to the likely occurrence of bacterial sulfate reduction in a closed or semiclosed environment (Goldhaber and Kaplan, 1980). This is because the rate of microbial sulfate reduction is faster for 32S 16 2⫺ O4 than for 34S 18O2⫺ as the energy required to break the 4 32 S–16O bond is lower than that for 34S–18O bond (Harrison and Thode, 1958; Lloyd, 1968; Mizutani and Rafter, 1969). Therefore, sulfate ions remaining in solution will progressively

Microbial sulfate reduction rates and sulfur and oxygen isotope fractionations

237

become enriched in heavy sulfur and oxygen isotopes as reduction proceeds. 3.3. Sulfate Reduction Rates Sulfate reduction rates were estimated from one dimensional diffusion-advection-reaction model of Berner (1964; 1980). The model describes the change in concentration of sulfate with time at a given depth x below the sediment surface as a function of diffusion, sediment accumulation and bacterial sulfate reduction in a steady-state system: D s共⭸ 2C/⭸ x 2兲 ⫺ ␻ 共⭸C/⭸ x兲 ⫺ f共 x兲 ⫽ 0,

(3)

where C is the sulfate concentration, D s is the diffusion coefficient of sulfate, ␻ is the sedimentation rate, x is the depth below the sediment-water interface, and f( x) is the depthdependent rate of sulfate reduction. The rate function f( x) can be calculated from the sulfate concentration profiles in Fig. 2 as follows: f共 x兲 ⫽ ae ⫺bx,

(4)

where a and b are constants. With the boundary conditions of C ⫽ C 0 for x ⫽ 0 and a finite value of C(C ⬁ ) for x going to infinity, Eqn. (3) has the following solution: C共 x兲 ⫽ 共C0 ⫺ C ⬁兲e ⫺bx ⫹ C ⬁,

(5)

共C 0 ⫺ C ⬁兲 ⫽ a/共D sb 2 ⫹ ␻ b兲.

(6)

where

Equation (5) predicts an exponential decrease of the sulfate concentration with depth approaching asymptotically the value of C ⬁ . In order to calculate the rate of sulfate reduction f( x), constants a and b are determined from Eqns. (5) and (6) above as follows: a ⫽ 共C ⫺ C ⬁兲/共D sb 2 ⫹ ␻ b兲,

(7)

where (C ⫺ C ⬁ ) and b are determined by fitting an exponential function to the observed sulfate profiles in the sediment cores (Fig. 4). For example, the best fit for the sulfate data in core 2647-4 is C ⫽ 4.5e ⫺0.059x ⫹ 24.5 ␮mol SO2⫺ cm⫺3. Thus, (C ⫺ 4 C ⬁ ) ⫽ 4.5 mol cm⫺3 and b ⫽ 0.059. Substituting these values into Eqn. (7) and taking an average sedimentation rate (␻) of 6 cm/1000 yr for the Gulf of Mexico slope sediments (Aharon, 1991) and a diffusion coefficient (D s ) of 100 cm2 yr⫺1 (Berner, 1978), we calculate a ⫽ 0.0043. The depthdependent rate of sulfate reduction for core 2647-4 is f( x) ⫽ 0.0043e ⫺0.059x ␮mol SO2⫺ cm⫺3 day⫺1. It follows that the 4 maximum sulfate reduction rate in the reference sediment is 0.0043 ␮mol SO2⫺ cm⫺3 day⫺1 at x ⫽ 0. Using the same 4 approach, the integrated reduction rates for the chemosynthetic sediments were calculated and the values are listed in Table 2. Although the rates calculated for the microbial sulfate reduction in seep sediments are highly variable, from 0.004 to 2.51 ␮mol SO2⫺ cm⫺3 day⫺1, and the number of core sediments 4 analyzed (n ⫽ 9) is too small to draw firm conclusions, an interesting pattern emerges. First, sulfate reduction rates increase progressively from reference (0.0043 ␮mol SO2⫺ cm⫺3 4

Fig. 4. Modelled downcore sulfate removal through bacterial sulfate reduction in reference core (A) and cores representative of oil seeps (B) and gas seeps (C) (concentration vs. core-depth profiles shown in Fig. 2). Note the coherence achieved between measured (solid circles) and modelled (crosses) data.

⫺3 day⫺1) to oil seeps (0.01 to 0.22 ␮mol SO2⫺ day⫺1) and 4 cm 2⫺ ⫺3 ⫺1 to gas seeps (0.27 to 2.51 ␮mol SO4 cm day ). Secondly, and most intriguing, sulfate reduction rates in gas seeps are higher by up to 260 times relative to the oil seeps. An explanation for the highly disparate sulfate reduction rates between seeps is given below.

4. DISCUSSION

4.1. Source of the Pore Fluids Knowledge of the pore fluid source/s is critical for a correct assessment of the sulfate starting point and for tracing its fate

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P. Aharon and B. Fu

Table 2. Estimates of sulfate reduction rates in oil and gas seep sediments, Gulf of Mexico. Sampling Areas

Core Number

a*

b*

Integrated Rate (␮mol SO2⫺ cm⫺3 day⫺1) 4

Reference

2647-4

0.004

0.059

0.0043

Oil seeps

2635-1 2635-2 2635-3 2639-2 2639-3 2639-4

0.010 0.038 0.096 0.074 0.220 0.101

0.037 0.144 0.391 0.110 0.179 0.140

0.0097 0.0377 0.0956 0.0737 0.2196 0.1013

Gas seeps

2647-1 2637-2 2647-3

2.512 1.492 0.267

0.602 0.435 0.185

2.5122 1.4924 0.2672

* a and b are constants used to calculate reduction rates from pore-fluid sulfate profiles (see text).

during the microbial reduction. Fu and Aharon (1998) have discerned three types of pore fluids associated with hydrocarbon seeps in the Gulf of Mexico on the basis of their chemical and isotopic fingerprints: (i) type I fluids having normal Gulf of Mexico deep water salinities (about 38‰) and originating as seawater trapped in the sediments whose chemical and isotope compositions were modified by microbial processes, authigenic carbonate precipitation, and gas hydrate sublimation; (ii) type II fluids which are highly saline (116‰ to 182‰ salinity range) and are derived by dissolution of subsurface salt diapirs during seawater convective circulation; and (iii) type III fluids whose salinities are intermediate between types I and II (52‰ to 155‰), are highly enriched in Ba, Ra, Sr, and Ca, and are derived from deep-seated formation waters advecting on the seafloor. Two lines of evidence argue the case that pore fluids studied here have a normal marine origin and pertain to type I fluids above. First, Cl concentrations in pore fluids from both oil and gas seeps (range 558 to 583 mM/L) are identical within analytical error to the Cl concentration in the overlying bottom water (Cl ⫽ 568 mM/L) and are practically invariable downcores (Table 1). Salinities measured by hand-held refractometer immediately after pore fluids recovery are in agreement with the Cl concentrations in the same pore fluids measured by ion chromatography (Table 1) and are identical to the salinity of the overlying Gulf of Mexico waters at 600 m depth (S ⫽ 38‰). Second, 87Sr/86Sr determinations of three representative pore fluids (2635-1-6; 2639-2-8 and 2647-2-8, Table 1) yield values (0.70916; 0.70917; 0.70917) which are indistinguishable from the 87Sr/86Sr ratio of the overlying bottom water (0.70917). It can be concluded, therefore, that the fluids trapped in the sediment originated as normal seawater and the low and highly variable SO4/Cl molar ratios listed in Table 1 (range 3.5 to 0.07) relative to the ambient seawater (5.2) are strictly the result of SO4 consumption during microbial sulfate reduction.

shallow, near shore, organic carbon-rich environments (up to 3 ␮mol SO2⫺ cm⫺3 day⫺1, Edenborn et al., 1987). Lin and 4 Morse (1991) measured sulfate reduction rates in sediments of the continental shelf and slope in the northern Gulf of Mexico and reported that the rates decrease exponentially with water depth (Fig. 5). The diminishing sulfate reduction rates with depth was attributed (op. cit.) primarily to the accompanying decrease of organic matter deposition which is controlled by the rate of sedimentation. In addition, at low sedimentation rates typical of deepwater, labile and metabolizable organic matter undergoes oxic and suboxic degradation in the surface

4.2. Sulfate Reduction Rate in Seeps

Fig. 5. Relationship between sulfate reduction rate and water depth in the Gulf of Mexico. Crosses represent data from normal marine sediments measured by Lin and Morse (1991). Data from reference (open circle), oil seeps (solid circles), and gas seeps (solid triangles) are listed in Table 2. Sulfate reduction rates in gas seeps overlain by methanotrophic mussels are anomalously high relative to marine sediments at comparable depth (see text).

Sulfate reduction rates in seep sediments, listed in Table 2, assume values which are intermediate between deep-sea or⫺3 ganic carbon-poor habitats (as low as 0.0002 ␮mol SO2⫺ 4 cm day⫺1, Bender and Heggie, 1984) and marine sediments from

Microbial sulfate reduction rates and sulfur and oxygen isotope fractionations

sediments for longer duration. Consequently, the abundance of labile organic matter buried in the sediment decreases resulting in low sulfate reduction rates (Lin and Morse, 1991). A comparison of sulfate reduction rates estimated in seeps with the rates measured by Lin and Morse (1991) may offer important insights on the systematics of microbial sulfate reduction as a function of the organic carbon substrate (Fig. 5). This is because our study sites are located on the same basin slope (northern Gulf of Mexico) as the sediments studied by Lin and Morse (1991) but with the notable difference that fossil hydrocarbons (i.e., either crude oil or methane gas) serve as the carbon substrate in the seep sediments whereas terrestrial/ marine sourced organic carbon serve as the substrate in the normal marine sediments. The feature standing out in this comparison is the observation that maximum sulfate reduction rates in gas seeps are anomalously high, by up to a factor of 10, relative to marine sediments (Lin and Morse, 1991) at comparable depths (Fig. 5). We notice, however, that the rates from reference and some oil seeps are somewhat lower than the ones measured by Lin and Morse (1991) at comparable depths, although intuitively we would have expected at least similar values. The observed discrepancy may be attributed either to (i) retardation of sulfate reduction in oil-saturated sediments as crude oil coats organic matter and prevents bacteria from metabolizing it (Ian Kaplan, personal communication), or (ii) different methodologies used for the determination of sulfate reduction rates. In this study we used a mathematical model to derive rates whereas the rates obtained by Lin and Morse (1991) were measured directly by using a 35S radiotracer technique. The discrepancies noted above may suggest that the two methods are not fully compatible. This contention is supported by the studies of Jørgensen (1978) who compared the two methods and reached a similar conclusion and by the observation that our oil-free reference sediment yields a sulfate reduction rate below the values obtained by Lin and Morse (1991) (Fig. 5). We contend, therefore, that sulfate reduction rates reported in the study are likely to represent minimum values of the true reduction rates. Two other questions need to be answered with respect to the data graphed in Figure 5, namely (i) why sulfate reduction rates, calculated by the same methodology, are generally higher in seep sediments than the reference sediment; and (ii) what causes the large difference observed between sediments hosting oil and gas seeps? A number of factors have been shown by Bernard and Westrich (1984) to control rates of bacterial sulfate reduction, including (1) temperature; (2) pressure; (3) concentration of dissolved sulfate in pore fluids; and (4) amount of reactive organic matter entering the sediment and the extent to which it can be metabolized. It is generally accepted that sulfate reduction rate increases with both temperature and pressure (Goldhaber and Kaplan, 1975; Jørgensen, 1977; Elsgaard et al., 1994). However, bottom temperatures at the sampling sites in this study are between 6 and 9°C and the water depth ranges only between 590 and 630 m (i.e., 59 to 63 atm). These small differences in temperature and hydrostatic pressure are therefore unlikely to have caused the observed variations in sulfate reduction rates. Dissolved sulfate concentration could affect the reduction rates but only when the concentration is less than 10% of the seawater value (Martens and Berner, 1977; Jørgensen, 1981;

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Bernard and Westrich, 1984). After the sulfate concentration drops to less than 10% of the seawater value (⬍3 mM/L), the reduction rate decreases with the sulfate depletion (Westrich, 1983). Most pore fluids considered here contain more than 10% of the seawater sulfate source. In addition, most pore fluids having a significant depletion of sulfate were found in gas seeps where the highest reduction rates are observed (Table 2). Therefore, sulfate concentration cannot be the primary controlling factor for the reduction rates in seeps. It is widely accepted that the amount and metabolic reactivity of the organic matter in sediments are the two most important factors controlling sulfate reduction rates (Goldhaber and Kaplan, 1975; Berner, 1978, 1980; Jørgensen, 1981; Westrich, 1983). Of these two variables, the metabolic reactivity of the organic matter is considered to be far more important than simply its amount (Westrich, 1983). Therefore, the higher sulfate reduction rates observed in seep sites relative to the reference can be attributed to the presence of a normal marine carbon substrate supplemented by additional seeping hydrocarbons and authochtonous dead chemosynthetic biomass. The excess carbon alone, however, cannot explain the large discrepancies observed between the gas and oil seeps. We propose that the difference in reduction rates is controlled by the types of hydrocarbons reaching the respective seep sites. Submersible observations indicate that gas hydrates at the threshold of stability (Brooks et al., 1984) and containing thermogenic methane are common at the coring sites where methanotrophic mussels abound whereas crude oil is the main form of hydrocarbon in sediments overlain by bacterial mats (Aharon et al., 1992; Sassen et al., 1993). Contrary to previous views (Atlas, 1981; Leahy and Colwell, 1990), recent laboratory experiments have demonstrated that hydrocarbons in crude oil form can be oxidized directly by some species of sulfate-reducing bacteria (Rueter et al., 1994). Sulfate reduction using crude oil in the bacterial mats was also inferred based on the modeling of ␦13C negative anomalies measured in the bottom waters (Aharon et al., 1992) and pore fluids (Fu, 1998). Because the molecule of methane is smaller and considerably simpler than that of refractory crude oil (Ferry, 1997), it is reasonable to suggest that bacteria prefers methane over crude oil. Thus, it is to be expected that sulfate reduction rates coupled with anaerobic methane oxidation in gas seeps tend to be higher than those coupled with anaerobic oxidation of crude oil in oil seeps. This interpretation is supported by the recent study of Devol and Ahmed (1981) which demonstrated that the sulfate reduction using methane substrate typically results in unusually high reduction rates. Alternatively, methanotrophs at the sediment surface may supply the readily metabolizable organic matter for sulfate reduction while the flow of methane maintains the reducing conditions in the sediments and renders the organic matter more accessible (Ian Kaplan, personal communication). 4.3. Sulfur and Oxygen Isotope Fractionations Bacteria are known to discriminate to various degrees against the heavier 34S and 18O isotopes (Harrison and Thode, 1958; Mizutani and Rafter, 1969). The sulfate/sulfide complementary curves in the seep sediments (Fig. 2) and the inverse relation observed between dissolved sulfate, ␦34S and ␦18O values (Fig. 6) are compatible with a sulfate reduction occur-

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Fig. 6. Relations between sulfur and oxygen isotope departures from marine sulfate values (␦34S ⫽ 20.3‰; ␦18O ⫽ 9.7‰) and the residual sulfate fraction in seep pore fluids. The nonlinear behavior of the data suggests variable isotope fractionations.

ring in a closed or semiclosed pore-fluid system. Hence the fractionation factors (␣) of S and O isotopes during sulfate reduction can be derived from a Rayleigh equation based on measurements of the isotope ratios and sulfate concentrations in the pore fluids (Broecker and Oversby, 1971): R t/R 0 ⫽ 共N t/N 0兲 ␣⫺1,

(8)

where R t and R 0 are the measured isotope ratios of sulfate at time t(N t ) and at the initiation of sulfate reduction (N 0 ); ␣ is the fractionation factor for S or O, respectively; (N t /N 0 ) is the fraction of sulfate ( f ) remaining in the pore fluids at any given time. Equation (8) can be rewritten as follows (Aharon, 1985): ⌬ ␦ ⫽ ␦ t ⫺ ␦ 0 ⫽ 10 3共 ␣ ⫺ 1兲ln f,

(9)

where ␦ t ⫽ ␦ 34S or ␦ 18O values of the residual SO2⫺ 4 fraction; ␦0 ⫽ ␦34S or ␦ 18O values of the initial sulfate before microbial sulfate reduction commenced (seawater sulfate: ␦34S ⫽ 20.3‰;

Fig. 7. Sulfur isotope enrichment factors (␧) relative to a marine value of ␦34S ⫽ 20.3‰ as a function of residual sulfate fraction left in the pore fluids after bacterial sulfate reduction. The plots are separated according to the type of seep [(A) nonseep reference; (B) oil seeps; (C) gas seeps].

␦18O ⫽ 9.7‰); 103(␣ ⫺ 1) is also known as the enrichment factor (␧) in permil. In a plot of ⌬␦ versus ln f, the slope representing the enrichment factor ␧ is invariable if the data points follow a straight line. Contrary to a straight line behavior, however, our data indicate that changes in SO4 during bacterial reduction in seep sediments are accompanied by nonlinear changes in the isotope compositions which vary substantially between seep sites (Fig. 6). Therefore pore fluid data in Figures 7 and 8 are separated according to the sediment type in order to derive the individual enrichment factors on the basis of their respective slopes. The enrichment factors (␧) of sulfur and oxygen isotopes are highest in pore fluids from the reference sediment (27.2‰ and 14.7‰) but substantially lower in oil seeps (18‰ and 6.2‰) and gas seeps (8.6‰ and 2.3‰). The sulfur isotope fractionations derived from the reference sediment (␣ ⫽ 1.027) and from oil seeps (␣ ⫽ 1.018) compare

Microbial sulfate reduction rates and sulfur and oxygen isotope fractionations

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during sulfate reduction in marine sediments have not been previously determined. Fractionation factors based on laboratory studies which range from 1.002 to 1.008 (Lloyd, 1968; Mizutani and Rafter, 1969) are compatible with our measured values of 1.006 and 1.002 in oil and gas seeps, respectively, but are substantially lower compared to the value of 1.015 derived in the reference sediment. The question of the factor(s) controlling the isotope fractionations during bacterial sulfate reduction needs to be addressed. These factors may be ambient, such as temperature and pressure, or may be intrinsic to the microbial communities, such as the rates of sulfate reduction. In this study, ambient factors can be excluded as the cause of the variable fractionations because reference and chemosynthetic sediments occur at similar water depths having similar bottom temperatures and pressures. This leaves the microbial factor as the primary suspect for controlling the isotope fractionations. Inverse relationships between sulfate reduction rates and sulfur and oxygen isotope fractionations is clearly established in Figure 9 suggesting that the former acts as the primary controlling factor. Therefore, the progressive decrease of sulfur and oxygen isotope fractionation factors from the reference sediment to oil seeps to gas seeps is likely caused by the increase of reduction rates. These observations from submarine seeps confirm the results of previous studies concerning the link between sulfur isotope fractionation and sulfate reduction rates (Jones et al., 1956; Harrison and Thode, 1958; Kaplan, 1962; Kaplan and Rittenberg, 1964; Kemp and Thode, 1968; Chambers et al., 1975; Habicht and Canfield, 1996, 1997) and also document the fact that a similar link exists with the oxygen isotopes albeit with differences in slopes (Fig. 9). 4.4. Mechanism of Sulfur and Oxygen Isotope Fractionations

Fig. 8. Sulfate-derived oxygen isotope enrichment factors (␧) relative to a marine value of ␦18O ⫽ 9.7‰ as a function of residual sulfate fraction left in the pore fluids after bacterial sulfate reduction. The plots are separated according to the type of seep [(A) nonseep reference; (B) oil seeps; (C) gas seeps].

favorably with isotope fractionation data derived from porefluid sulfate profiles (Table 3) and laboratory experiments using Desulfovibrio desulfuricans which generally centers around ␣ ⫽ 1.025 (Kaplan, 1962; Kaplan and Rittenberg, 1964; Kemp and Thode, 1968). In contrast, the sulfur isotope fractionation derived from gas seeps is much lower (␣ ⫽ 1.009) than the preceding values. It must be stressed that sulfur isotope fractionations obtained on the basis of differences between presumed contemporaneous sedimentary sulfide and sulfate minerals typically exhibit a wide range (␣ ⫽ 1.016 to 1.070) (Kaplan, 1962; Hartman and Nielsen, 1969; Goldhaber and Kaplan, 1980; Habicht and Canfield, 1997) which has been attributed by some to the effect of highly variable sulfate reduction rates (Rees, 1973; Goldhaber and Kaplan, 1980) and by others to the effect of sulfide reoxidation (Jørgensen, 1990; Canfield and Thamdrup, 1994). To the best of our knowledge, oxygen isotope fractionations

The ␦34S and ␦18O compositions of SO4 consumed during bacterial sulfate reduction may be controlled either by (i) decoupled sulfur and oxygen isotope fractionations involving kinetic isotope exchange for sulfur but equilibrium isotope exchange of oxygen with ambient H2O, or (ii) coupled sulfur and oxygen isotope fractionations resulting from kinetic effects governed by microbial reduction pathways. The problem of distinguishing between the two controlling mechanisms of isotope fractionations above is still unresolved, among others, because of the rarity of combined ␦34S and ␦18O determinations of the residual sulfate. Bacterial sulfate reduction is a complex biochemical process which entails the formation of sulfate-enzyme complexes as well as sulfite as intermediate products between external sulfate and sulfide (Harrison and Thode, 1958; Rees, 1973): (1) (2) (3) SO42⫺(ext) 3 SO42⫺ enzyme 3 SO32⫺ 3 S2⫺. In this model, dissolved sulfate from the ambient environment is assimilated by the bacterial cells. The internal sulfate then reacts with the ATP (Adenosine TriPhosphate) to form APS (Adenosine-5⬘-PhosphatoSulfate) which thereafter is reduced directly to sulfite (Habicht and Canfield, 1997). The final step of this pathway involves the reduction of sulfite to sulfide. Rees (1973) has proposed that, in the absence of rate limi-

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Table 3. Comparison of sulfur isotope fractionations calculated based on pore-fluid sulfate profiles in this study and previous studies. Location Kiel Bay, Baltic Sea

... ... ... Pescadero Basin, Gulf of California Carmen Basin, Gulf of California Site 27 (JOIDES) Southwest Atlantic Ocean Santa Barbara Basin ... West Coast of Baja California ... ... Gulf of Mexico ... ...

Sediment Type

Water Depth (m)

Core Depth (m)

Isotope Fractionation Factor (␣)

Reference*

Shelf sediments ... ... ...

19 26 24 22 3361 2772

Reference Oil seeps Gas seeps

591 624 591

0.4 0.3 0.4 0.3 1.6 2 249 ⬎0.15 ⬍0.15 ⬎0.4 ⬍0.4 0.7 0.2 0.2 0.3

1.017 1.030 1.026 1.028 1.036 1.033 1.026 1.042 1.025 1.034 1.019 1.032 1.026 1.018 1.009

(1) and (2) ... ... ... (3) (3) (4) (4) ... (4) ... ... (5) ... ...

* (1) Hartmann and Nielsen (1969); (2) Chambers and Trudinger (1978); (3) Goldhaber and Kaplan (1980); (4) Kaplan (1983); (5) this study.

tation, the overall isotope fractionation should be the sum of the individual kinetic isotope effects occurring in the reaction steps above, (i.e., (i) about 3‰ 34S enrichment of the internal, relative to the external, sulfate in step 1, and (ii) up to 47‰ depletion in 34S for the reduction in steps 2 and 3 due to the splitting of S–O bonds (Harrison, 1957; Harrison and Thode, 1958; Rees, 1973). Concerning the oxygen isotopes, there are two sites where sulfate could exchange its isotopes with waterderived oxygen (Mizutani and Rafter, 1973), either during rapid equilibration of SO4-enzyme complex (equilibrium fractionation) or during re-oxidation of SO3 back to sulfate thus implicating the breakage of H–O bond (kinetic fractionation). Fritz et al. (1989) argued that isotope exchange between SO4enzyme complexes and water dominates the bacterial sulfate reduction on account of the following evidence: (i) isotopic

Fig. 9. Inverse relations between bacterial sulfate reduction rates and isotope enrichment factors in seeps. Both sulfur and oxygen–sulfate isotope fractionations decrease as reduction rates increase.

enrichment of the residual sulfate in 34S unaccompanied by a corresponding 18O enrichment, and (ii) fractionation data approaching the temperature-dependent thermodynamic isotope effects predicted for the SO4–H2O system (Lloyd, 1968; Mizutani and Rafter, 1969). Whether or not kinetic effects or “steady state” isotope exchange governs the fractionations in seep sediments is tested in a cross plot of ␦18O/␦34S data in Figure 10. Were the sulfate ␦18O data controlled by thermodynamic isotope effects and equilibrium exchange with the large excess of O in the pore fluids H2O (␦18Ow ⫽ 1.5 to 2‰ SMOW), a “plateau” at around

Fig. 10. ␦18O vs ␦34S sulfate data in pore fluids from reference sediment (empty circle), oil seeps (solid circles), and gas seeps (solid triangles) display a positive relation suggesting that kinetic isotope effects control both sulfur and oxygen isotope fractionations during bacterial sulfate reduction. The predicted range of equilibrium isotope fractionations between SO4–H2O at the ambient conditions of seeps (␦18O ⫽ 1.5 to 2.0‰ SMOW; t ⫽ 6 to 9°C) are indicated by dotted patterns 1 and 2 estimated from the experimental equations of Lloyd (1968) and Mizutani and Rafter (1969), respectively.

Microbial sulfate reduction rates and sulfur and oxygen isotope fractionations

the equilibrium value of between 35‰ (Lloyd, 1968) to 30.5‰ (Mizutani and Rafter, 1969) would be expected at cold ambient water temperatures of 6°C to 9°C measured at our seep sites. Instead, our ␦18O data are variable (ranging between 9 to 24‰ SMOW) and are exhibiting a positive linear relationship with the paired ␦34S values (Fig. 10). Our observations thus contradict the ones reported by Fritz et al. (1989) and argue in favor of kinetic isotope effects controlling both sulfur and oxygen isotope fractionations during sulfate reduction in seep sediments. The most likely site for both sulfur and oxygen isotope fractionations is step 2 above which involves the cleaving of the SO4 molecule and partial reoxidation of sulfite back to sulfate. The reoxidation of sulfite to sulfate may explain the observed ␦18O variations of the residual sulfate (Fig. 10). We point out that Fritz et al. (1989) interpretation of their ␦34S and ␦18O data is considerably weakened by experimental difficulties involved in controlling a fixed initial SO4 concentrations, and failure to avoid 18O enrichments in the water due to evaporation. On the basis of ␦34S data alone, Habicht and Canfield (1997) have proposed that there is a link between sulfate supply, sulfate reduction rates, and degree of isotope fractionation. According to this model, higher fractionations (⬎22‰) will occur under conditions of unlimited SO4 supply and low sulfate reduction rates whereas limiting SO4 supply and/or high sulfate reduction rates will lead to lower fractionations (⬍22‰). Our seep data is in agreement with Habicht and Canfield (1997) model discussed above. First, an inverse relation exists between the estimated sulfate reduction rates and the isotope fractionations shown in Figure 9. In the reference sediment the isotope fractionation factor is about ␣ ⫽ 1.027. We attribute this high fractionation value to a relatively slow sulfate reduction rate which results in the reduction of sulfate to sulfite (step 2) to be the rate-limiting. In contrast, the relatively lower sulfur isotope fractionations observed in oil seeps (␣ ⫽ 1.018) and gas seeps (␣ ⫽ 1.009) are attributed to the fast reduction of sulfate and the lower sulfate concentrations in these seep sediments. Under these circumstances, step 1 may become partially rate limiting and may alternate with steps 2 and 3 resulting in a sulfur isotope effect ranging between 1.000 and 1.022.

4.5. Implications of the Seeps Isotope Data Hydrocarbon seeps serve as a convenient natural laboratory for investigating sulfur and oxygen isotope fractionations resulting from microbial sulfate reduction. This is because (i) sulfate reduction occurs close to the sediment/water interface and it is completed within the uppermost 50 cm; (ii) rates vary substantially between sites according to the type of hydrocarbon substrate; (iii) ambient conditions are practically invariable in deepwater thus minimizing the factors controlling the observed variable fractionations, and (iv) the chemosynthetic biota inhabiting the seeps are diagnostic of the flux rates of sulfide production in the underlying sediment, i.e., Beggiatoa mats and Bathymodiolus mussels being associated with slow and fast fluxes, respectively. Two implications become manifest in light of the data presented above.

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Fig. 11. Schematic model indicating that both equilibrium (partial or complete) fractionations between SO4–H2O, accelerated by high temperatures (t ⬎ 50°C), and kinetic fractionations caused by variable sulfate reduction rates could account for changes in ␦34S/␦18O ratios of residual sulfate. In the case of deepwater seeps, the high temperature factor can be ruled out leaving the variable rates as the primary cause for the observed isotope fractionations (see text).

4.5.1. Variable ␦34S/␦18O ratios in sulfate Sulfate isotope data plotted in ␦ 34S/␦ 18O coordinates may offer important information concerning the origin of the sulfate and its geochemical history (Fig. 11). A number of authors have proposed on the basis of laboratory studies of sulfate solutes (Rafter and Mizutani, 1967; Mizutani and Rafter, 1969) and sedimentary barites (Sakai, 1971) that isotope data initiating at seawater values and straddling along a line with a slope of 4 (i.e., ␦34S ⫽ 4 ␦18O) are diagnostic of a seawater sulfate source isotopically modified during microbial sulfate reduction. Claypool et al. (1980) justified the slope of 4 on account of an average enrichment factor for sulfur which is 4 times larger than the oxygen isotope enrichment factor. Documentation of ␦34S/␦18O ratios deviating from 4 either remained unexplained (e.g., Cecile et al., 1983) or were attributed to factors independent of microbial metabolism (e.g., Aquilina et al., 1997). Both thermodynamic and kinetic factors can cause a decrease of the isotope enrichment ratios, including (Fig. 11): (i) partial and/or complete re-equilibration of the SO4-oxygen with H2O-sourced oxygen; and (ii) variable ␧34S/␧18O ratios controlled by the rates of sulfate reduction (Figs. 7 and 8). The former implicates temperatures substantially higher than ambient in order to speed up the sluggish equilibration rate between sulfate and water (Zak et al., 1980). Plots of pore fluid data in ␦34S and ␦18O coordinates (Fig. 12) reveal enrichment factor ratios ranging from ␧34S/␧18O ⫽ 1.4 in the reference sediment to 2.8 in oil seeps and 3.5 in gas seeps. In the absence of neither high temperatures, nor temperature variability, a thermodynamic factor can be dismissed in deepwater seeps. It must therefore be concluded that kinetic factors related to sulfate reduction rates control the isotope

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sulfate reduction in a sulfate-rich ocean (Ohmoto and Felder, 1987; Ohmoto et al., 1993), and (ii) sulfide formation in a sulfate-poor ocean (Cameron, 1982; Schidlowski et al., 1983; Hattori et al., 1983). The first argument is based on the observation that 34S depletion in sulfide decreases as the reduction rate increases (Kaplan and Rittenberg, 1964; Chambers et al., 1975; Goldhaber and Kaplan, 1975; Ohmoto et al., 1993). Accordingly, the small isotopic differences between sedimentary sulfides and coeval sulfate of the Archean oceans were attributed to rapid sulfate reduction in a sulfate-rich ocean (Schidlowski, 1979; Ohmoto et al., 1993). Viewed this way, sulfate reduction rates of between 27 to 270 ␮mol cm⫺3 day⫺1 have been speculated (Ohmoto et al., 1993). Alternatively, the small 34S depletion in the Archean sulfides was interpreted as having a mantle origin on account of the similarity of the sulfides ␦34S values to that of meteorites and mantle-derived igneous rocks (Monster et al., 1979; Schidlowski, 1979; Cameron, 1982; Hattori et al., 1983; Lambert and Donnelly, 1990). The choice between these two scenarios, and consequently the answer to the question of high or low levels of sulfate in Archean oceans, rests on the ability of bacteria to reduce sulfate at high rates with minimal isotope fractionation. Our seep data (Fig. 11) provide support for the argument relevant to the Archean oceans that rapid sulfate reduction can be associated with lesser isotope fractionation as enrichment factors as low as 8.6‰ and 2.3‰ for sulfur and oxygen isotopes were documented in sediments during anaerobic methane oxidation.

5. CONCLUSIONS

Fig. 12. Plots indicating variable ␦34S/␦18O ratios in (A) reference sediment, (B) oil seeps, and (C) gas seeps.

enrichment ratios, i.e., slow rates are associated with low ratios while faster rates are associated with progressively higher ratios. Aquilina et al. (1997) reported recently a ␦34S/␦18O ratio of 2 in barites from seeps at the convergent margin off Peru. The low ␦34S/␦18O ratio in the Peruvian barites was explained by these authors (op. cit.) to be caused by an 18O enrichment in SO4 resulting from equilibrium isotope exchange between SO4 and seawater fostered by high temperature (⬎100°C). The water temperature data, however, does not support any significant heat anomalies in their study area. In light of the findings from the Gulf of Mexico seeps, the lower ␦34S/␦18O ratio of these barites can be conceivably attributed to slow microbial sulfate reduction prior to the formation of the Peruvian barites. 4.5.2. Origin of Archean sulfides Sedimentary sulfides of Archean age typically display small depletions in 34S values (⬃0 ⫾ 5‰) relative to contemporaneous seawater sulfate (Schidlowski, 1979; Schidlowski et al., 1983). Two distinct arguments have been developed in order to explain these small 34S depletions in Archaen sulfides: (i) rapid

(1) Hydrocarbon seeps occurring at 590 to 630 m depths in the Gulf of Mexico serve as a convenient laboratory to explore sulfur and oxygen isotope systematics of sulfate in pore fluids because of intense microbial sulfate reduction, unusual at bathyal depths, coupled with practically invariable ambient temperatures (6°C to 9°C) and a stable source of seawaterderived sulfate. (2) Pore fluids acquired from oil and gas seeps overlain by chemosynthetic fauna of bacterial mats and mussel beds, respectively, are highly depleted in SO4 (down to 0.3 mM/L) and enriched in H2S (up to 21 mM/L) relative to bottom waters (SO4 ⫽ 29 mM/L). About 31% of the sulfide produced by microbial sulfate reduction is removed by thiotrophs and precipitation of sulfides. (3) Sulfate reduction rates are substantially higher in oil and ⫺3 gas seeps, up to 0.22 and 2.5 ␮mol SO2⫺ day⫺1, respec4 cm tively, compared to a nonseep reference sediment at comparable depth (0.0043 ␮mol SO2⫺ cm⫺3 day⫺1). The major factor 4 controlling the variable reduction rates in seeps is the type of hydrocarbon-derived carbon mineralized by the sulfate reduction bacteria. (4) Sulfur and oxygen isotope fractionations in seeps range from ␣ ⫽ 1.009 to 1.026 and from ␣ ⫽ 1.002 to 1.015 for sulfur and oxygen, respectively. These isotope fractionations are inversely related to the reduction rates. (5) Contrary to previous reports claiming that ␦ 18O compositions of sulfate are determined by equilibrium isotope exchange between sulfate and water, this study indicates that both

Microbial sulfate reduction rates and sulfur and oxygen isotope fractionations

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