Sulfur-33 Constraints on the Origin of Secondary Pyrite in Altered Oceanic Basement

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Geochimica et Cosmochimica Acta 87 (2012) 323–340 www.elsevier.com/locate/gca

Sulfur-33 constraints on the origin of secondary pyrite in altered oceanic basement

Shuhei Ono a,⇑, Nicole S. Keller a,b,1, Olivier Rouxel c,d, Jeﬀrey C. Alt e

a Department of Earth, Atmospheric, and Planetary Sciences, Massachusetts Institute of Technology, 77 Massachusetts Avenue, Cambridge, MA 02139, USA b Geology and Geophysics Department, Woods Hole Oceanographic Institution, Woods Hole, MA 02543, USA c IFREMER, Centre de Brest and European Institute for Marine Studies, Place Nicolas Copernic, 29280 Plouzane´, France d Marine Chemistry and Geochemistry Department, Woods Hole Oceanographic Institution, Woods Hole, MA 02543, USA e Department of Earth and Environmental Sciences, University of Michigan, Ann Arbor, MI 48109, USA

Received 17 December 2011; accepted in revised form 4 April 2012; available online 11 April 2012

Abstract

Low temperature alteration of oceanic basement rocks is characterized by net gain of sulfur, which commonly yields low d34S values, suggesting involvement of microbial sulfate reduction. In order to test whether secondary sulfide minerals are consistent with a biogenic source, we apply high precision multiple sulfur isotope analysis to bulk rock sulfide and pyrite isolates from two contrasting types of altered oceanic basement rocks, namely serpentinized peridotites and altered basalts. Samples from two peridotite sites (Iberian Margin and Hess Deep) and from a basalt site on the eastern flank of the Juan de Fuca Ridge yield overlap- ping d34S values ranging from 0& to 44&. In contrast, sulfides in the basalt site are characterized by relatively low D33S values ranging from 0.06& to 0.04&, compared to those from peridotite sites (0.00& to 0.16&). The observed D33S signal is significant considering the analytical precision of 0.014& (2r). We present a batch reaction model that uses observed d34SandD33Srelation- ships to quantify the effect of closed system processes and constrain the isotope enrichment factor intrinsic to sulfate reduction. The estimated enrichment factors as large as 61& and 53&, for peridotite and basalt sites respectively, suggest the involvement of microbial sulfate reduction. The relatively high D33S values in the peridotite sites are due to sulfate reduction in a closed system environment, whereas negative D33S values in the basalt site reflect open system sulfate reduction. A larger extent of sulfate reduction during alteration of peridotite to serpentinite is consistent with its higher H2 production capacity compared to basalt alteration, and further supports in-situ microbial sulfate reduction coupled with H2 production during serpentinization reactions. Ó 2012 Elsevier Ltd. All rights reserved.

1. INTRODUCTION anic crust likely contributes to the long-term chemical evolution of mantle and Earth surface reservoirs and influences Seawater circulation through mid-ocean ridge systems the redox, C and S budgets (e.g., Lećuyer and Ricard, 1999; drives water–rock reactions that exert important controls Hayes and Waldbauer, 2006). An increasing number of on the chemical evolution of the oceanic crust and seawater studies suggest that the alteration of oceanic crust not only (Elderfield and Schultz, 1996). Subduction of altered oce- leads to critical chemical exchange at the crust-seawater interface, but that it may also provide a vast habitat for the deep subsurface biosphere (Cowen et al., 2003; Orcutt ⇑ Corresponding author. Address: Earth, Atmospheric, and et al., 2011). It is hypothesized that redox reactions associ- Planetary Sciences, Massachusetts Institute of Technology, E25- ated with water–rock interaction during alteration can 641, 77 Massachusetts Avenue, Cambridge, MA 02139-4307, USA. generate some limited quantity of energy to support micro- Tel.: +1 617 253 0474; fax: +1 617 253 8630. E-mail address: [email protected] (S. Ono). bial life in the deep subsurface (e.g., Bach and Edwards, 1 Present address: Institute of Earth Sciences, University of 2003). However, evidence for microbial life in oceanic base- Iceland, Sturlugata 7, Askja, 101 Reykjavik, Iceland. ment rocks remains limited or inconclusive.

Microbial alteration of basaltic glass from the upper oce- basaltic basement rocks is commonly characterized by per- anic crust has received significant attention, in part because vasive oxidation, where ferrous minerals (olivine and the alteration effects are clearly observable with traditional pyroxene) are oxidized to ferric minerals (e.g., celadonite petrographic techniques (microscopy, SEM, microprobe; (Fe3+mica), goethite). Oxidation of ferrous iron largely oc- e.g., Fisk et al., 1998; Furnes and Staudigel, 1999; Banerjee curs with dissolved O2 as the oxidant: and Muehlenbachs, 2003). Geochemical studies of sulfur 4FeO þ O þ 2H O ! 4FeOðOHÞð2Þ isotope ratios in secondary pyrite and carbon isotope ratios 2 2 in vein carbonate, in particular, have also been used to test where FeO and FeO(OH) represent the Fe2+ component of microbial activity in peridotitic (e.g., Alt and Shanks, 1998; ferrous minerals and the Fe3+ component of ferric miner- Alt et al., 2007; Delacour et al., 2008) and basaltic basement als), respectively (Bach and Edwards, 2003). rocks (Krouse et al., 1977; Andrews, 1979; Rouxel et al., Alternatively, Fe2+ hydration and oxidation may result 2008; Alt and Shanks, 2011). However, a protracted history in the production of H2 (e.g., Stevens and McKinley, of water–rock interaction from initially hydrothermal to 1995; Bach and Edwards, 2003): subsequent low temperature alteration in the ridge flank re- 2FeO þ 4H2O ! 2FeOðOHÞþH2: ð3Þ charge zone, can lead to several generations of sulfide min- as shown experimentally (Stevens and McKinley, 1995). erals of multiple origins, and thus complicates the Significance of this reaction, however, has been debated interpretation of the measured isotopic signatures. since a later study demonstrated that the H production Serpentinization reactions and associated H production 2 2 from reaction (3) is only transitional (Anderson et al., are of particular interest in the context of microbial life in 1998). Provided that the reaction (3) does occur, basaltic oceanic basement rocks, as evidenced by the unique micro- basement rocks, which comprise the vast majority of sub- bial community found at the Lost City hydrothermal vent marine basement rocks exposed to seawater fluids, may field, which is dominated by an Archaea, Methanosarcinales support lithoautotrophic microbiology based on H (Bach with minor species including bacteria of close similarity to 2 and Edwards, 2003). sulfur-and methane-oxidizing bacteria (Schrenk et al., Regardless of putative H production in basaltic sys- 2004; Kelley et al., 2005). Similar microbial metabolisms 2 tems, culture-dependent and independent studies support have been inferred for subseafloor serpentinization environ- activity (or survival) of a wide variety of microbes including ments, but this has not yet been directly tested due to the methanogens and sulfate reducers, in both terrestrial (Ste- difficulty of sampling subsurface basement fluids. vens and McKinley, 1995) and seafloor basalt aquifers The serpentinization reaction produces H from the oxi- 2 (Cowen et al., 2003). Having set potential contamination dation of Fe(II) in olivine to Fe(III) in magnetite An ideal . aside, the question remains as to whether the microbial reaction can be written as (e.g., Alt and Shanks, 1998; communities in these basaltic aquifers represent lithoauto- McCollom and Bach, 2009): trophy fueled by H or heterotrophy supported by the flux Mg Fe SiO 1:37H O 0:5Mg Si O OH 2 1:8 0:2 4 þ 2 ! 3 2 5ð Þ4 of organic material from overlying sediments (Anderson olivine Serpentine et al., 1998; Cowen et al., 2003). The low 14C content of dis- þ 0:3MgðOHÞ2 þ 0:0067Fe3O4 þ0:067H2 ð1Þ brucite magnetite solved organic carbon in basement fluids sampled from Serpentinization proceeds over a range of temperatures Juan de Fuca Ridge suggests a chemosynthetic microbial from 320 °C to below 150 °C, and is accompanied by a community at this site (McCarthy et al., 2010). change in the stable sulfide mineral assemblage, (e.g., Klein Microbial sulfate reduction (MSR) catalyzes the reduc- and Bach, 2009; McCollom and Bach, 2009). Late stage ser- tion of sulfate coupled with H2, and the product hydrogen pentinization (and steatitization) is characterized by an in- sulfide can react with iron to precipitate sulfide minerals. crease in sulfur content, with a sulfide mineral assemblage MSR is known to produce large isotope effects (e.g., Kap- of millerite (NiS)-pyrite (FeS2)-polydymite (Ni3S4) associ- lan and Rittenberg, 1964; Canfield, 2001; Farquhar et al., ated with an increase in oxygen fugacity (e.g., Alt and 2003; Sim et al., 2011a,b), thus the study of sulfur isotope Shanks, 1998; Alt et al., 2007; Klein and Bach, 2009). systematics of secondary sulfide minerals can provide a crit- In contrast to peridotite serpentinization reactions, H2 ical tool to assess putative microbial activity in the oceanic production during low temperature aqueous alteration of crust. Previous studies showed that basaltic basement rocks basalt is still subject to debate (Stevens and McKinley, commonly yield 34S-depleted pyrites with sulfur isotope ra- 1995, 2000; Anderson et al., 1998; Bach and Edwards, tios as low as 45& with respect to VCDT (Krouse et al., 2003). Although high temperature (>250 °C) water-basalt 1977; Andrews, 1979; Rouxel et al., 2008; Alt and Shanks, 34 interaction has been shown to produce H2 and to quantita- 2011). Although these highly S-depleted sulfides suggest tively reduce SO4 to H2S (e.g., Shanks et al., 1981), the rate involvement of microbial sulfate reduction, the occurrence of water–rock reactions is significantly slower at lower tem- of inorganic pathways cannot be excluded. Andrews peratures. Low temperature alteration2 (sulfates and sulfides and sulfate may produce the highly S-depleted sulfide (Ohmoto and Lasaga, 1982). This temperature is close to the upper minerals. The disproportionation model, however, would temperature limit for life, currently demonstrated to be at 122 °C predict net loss of sulfur from the basement rocks. This (Takai et al., 2008). contradicts other geochemical studies that show local S. Ono et al. / Geochimica et Cosmochimica Acta 87 (2012) 323–340 325

60°N

1301 897 40°

20° 801 1268

0° 895 20°

40°

60S°

120° 150° 180° 210° 240° 270° 300° 330° 0°3°0 ° 60 90°E

Fig. 1. Location of the Ocean Drilling Program sites examined in this study (1301, 895, 897 and 1268) and in Rouxel et al. (2008 – Site 801). Two basalt sites (801 and 1301) and three peridotite sites (895, 897, and 1268) are shown in circles and squares, respectively. uptake of sulfur in basaltic basement (e.g., Alt, 2004; Rou- et al., 1998; Wheat et al., 2002) as well as directly in the drill xel et al., 2008; Alt and Shanks, 2011). holes at Site 1026, situated 1–2 km to the north (Cowen In addition to conventional 34S/32S ratio analysis, high- et al., 2003) using CORK (Circulation Obviation Retrofit precision measurements of 33S have been shown to provide Kit) instrumentation. Source seawater is inferred to enter additional constraints on the source(s) of sulfur in seafloor the basement 50 km to the SSW at the Grizzly Bare outcrop hydrothermal deposits (Ono et al., 2007; Peters et al., 2010). on the same basement ridge, and fluids flow generally These studies show that biogenic sulfide may be distin- northward through the basement of the buried ridge guished from inorganic sulfide sources by 33S even when (Wheat et al., 2000, 2002; Fisher et al., 2003). the d34S values of the samples are inconclusive about their Hole 1301B penetrates 265 m of sediment and 318 m origin. The goal of this study is to test whether high-preci- into the basement. The upper 86 m of basement were cased sion 33S analysis can be used to better constrain the origin and not cored because of the brecciated and rubbly nature of sulfur in oceanic basement rocks. We present a triple sul- of the rocks. Sediments include 216 m of turbidite overly- fur isotope (32S/33S/34S) model based on a simple batch ing hemipelagic clays. The basement consists of dominantly reaction which we use to test the sensitivity of the measured aphyric to highly phyric pillow basalt, minor massive basalt data on geochemical parameters (e.g., isotope fractionation and two 1 m intervals of basalt-hyaloclastite breccia, all factors, the extent of sulfate reduction, water–rock ratios), with N-MORB compositions (Fisher et al., 2005). The con- and to explain different 33S systematics found between sam- centration of sulfate in sediment pore waters decreases from ples from three peridotite sites (Hess Deep, ODP Site 895; the seawater value at the surface (28 mM) to near zero at Iberian Margin, ODP Site 897; Mid-Atlantic Ridge 15 57 mbsf as the result of sulfate reduction within the sedi- °N, IODP Site 1268) and one basalt site (Juan de Fuca, ment cover (Fisher et al., 2005). The sulfate concentration IODP Site 1301). In order to expand the dataset for basaltic remains near zero down to 125 mbsf then increases with environments we include in our discussion data from Site depth, to 16 mM in the lowermost sediments. 801, previously published by Rouxel et al. (2008) (Fig. 1). Basement fluids sampled at the Baby Bare warm springs have sulfate concentrations similar to that of the lowermost 2. MATERIALS sediments (18 mM; Wheat et al., 2000). The low sulfate (compared to seawater) content of basement fluids has been 2.1. Basaltic basement site, IODP Site 1301 interpreted to be the result of sulfate diffusion from the basement fluids into the overlying sediments (Elderfield Hole U1301B is located in 3.5 Ma crust on the eastern et al., 1999). Small-subunit rRNA genes cloned from sam- flank of the Juan de Fuca Ridge. This site lies along a bur- ples from the basement fluid at nearby (1 km) hole ied basement ridge paralleling the spreading axis (Fig. 1). 1026B showed that sulfate reducing microbes (Bacteria as Two basement exposures lie along this ridge: Baby Bare well as that related to Archaeoglogus) are common, suggest- 6 km to the SSW, and Mama Bare 8 km to the NNE. ing that biogenic sulfate reduction may occur within the Basement fluids in the area were sampled at warm springs basement. However, Wheat et al. (2004) showed a signifi- (65 °C) at various Baby Bare vent sites (e.g., Mottl cant sediment component in the silica contents of the 326 S. Ono et al. / Geochimica et Cosmochimica Acta 87 (2012) 323–340 borehole fluids, either from a poor seal of the casing in the 97 m of basement, while hole 897D cored 694 m of sedi- basement or via diffusion from sediment to basement, ment and 153 m into the basement (Sawyer et al., 1994). leaving open the possibility that the sampled microbial An additional two samples were taken from Site 637 (situ- rRNA was originally present in the sediment rather than ated approximately 150 km north of Site 897), where 212 m in the basement. of sediment overlie serpentinite, which could be recovered At the spreading axis, the basement cools rapidly to from 212 to 285.6 mbsf (Boillot et al., 1987). Tectonic ambient seawater temperatures after its emplacement but unroofing and uplift at about 125 Ma resulted in emplace- warms up again as it becomes buried under a non-perme- ment of serpentinite at the surface. The samples from this able sediment cover. At Juan de Fuca, the basement fluid site consist of chromium reducible (i.e., pyrite) sulfur ex- temperature increases from 16 °Cat22kmto64°Cat tracted from bulk rock samples (Table 5). 103 km from the ridge axis (Davis and Becker, 2002). The basement is locally isothermal due to vigorous fluid convec- 2.2.2. Hess Deep, ODP Site 895 tion within highly permeable basements (Fisher, 2005). The Hess Deep is a young (0.5–1 Ma) lower crust and upper low temperature alteration of basalt recovered from Site mantle assemblage that formed at the East Pacific Rise, and 1301 is characterized by Fe-oxyhydroxide-bearing oxida- was exposed by Cocos-Nazca rifting (Alt and Shanks, tion halos along fractures, similar to basalt alteration in 1998). Hole 895A and 895B penetrated 10–17 m of perido- other sites (e.g., Andrews, 1979; Alt et al., 1989; Alt, tite and gabbro, and Hole 895C penetrated 37.9 m of harz- 2004). These cm-sized dark (black) to brownish alteration burgite and dunite with minor gabbroic and basaltic dikes. halos contain Fe-oxyhydroxides, celadonite (Fe3+ mica), Hole 895D cored 94 m of mainly harzburgite and lesser du- and saponite (Fe2+ smectite). The halos may be zoned, with nite, with gabbroic and basaltic veins and dikes. Hole 895E, Fe-oxyhydroxides closest to the halo, then a celadonite-rich 300 m to the north, penetrated 87.6 m of mostly dunite with zone, and finally the saponite-rich host rock. Secondary layers of harzburgite and lesser troctolite and gabbroic pyrite is concentrated at a “pyrite front” which marks the rocks. Serpentinization at the Hess Deep is characterized boundary between the alteration halo and the host rock by relatively high temperature (>200 °C), low water/rock (Andrews, 1979; Alt and Shanks, 2011). These observations ratios, and low-sulfur mineral assemblage including NiFe suggest a relationship between Fe oxidation and pyrite for- alloy (Alt and Shanks, 1998). The samples from this site mation. Rock samples were divided into three rock types consist of chromium reducible (i.e., pyrite) sulfur extracted (grey host rock, black halo, and pyrite front) and sulfide from bulk rock samples (Table 5). was extracted from each of the various types (Table 4). The samples from this site consist of chromium reducible 2.2.3. Mid Atlantic Ridge, ODP Site 1268 sulfur (CRS) and acid-volatile sulfur (AVS) extracted from Site 1268 is located at 14°500N, 15 km northwest of the bulk rocks, as well as pyrite separates. Logatchev vent site (Bach et al., 2006). Hole 1268A penetrated 147.6 m into completely altered harzburgite and du- 2.2. Peridotite basement sites nite, with local mylonitic shear zones, and highly altered late-magmatic dikes. Peridotites were exposed by detach- The conventional (34S/32S) sulfur isotope geochemistry ment faulting and were altered in two general stages: initial of the three peridotite sites investigated in this study was de- serpentinization resulted in formation of serpentine + mag- scribed in detail in previous studies (Alt and Shanks, 1998; netite ± pyrite in the rocks, and a later stage of talc alter- Alt et al., 2007). Subsets of the Ag2S produced as a part of ation resulted in replacement of serpentine by talc. The these previous studies were used in the present study and samples from this site consist of CRS and AVS (Table 5). analyzed for multiple sulfur isotopes. The geology and geo- Primary mantle sulfides are not preserved in serpentine chemistry of the sites are briefly described below. rocks from Hole 1268A (Alt et al., 2007). Secondary sulfide minerals occur along serpentinite and sulfide-bearing vein- 2.2.1. Iberian Margin, ODP Site 897 lets and gabbroic dikes. Sulfide minerals in these veins are The drill holes on the Iberian Margin are located on a composed of millerite (NiS), which is altered to polydymite peridotite ridge at the seaward edge of the ocean-continent (Ni3S4) or violarite (FeNi2S4) and pyrite (FeS2). Veins of transitional crust of the Iberian Margin (ODP Site 897; pyrite and talc are present, but rocks affected by the later Fig. 1). The peridotite is intensively serpentinized with more stages of talc replacement generally lack sulfide minerals, than 90% of the primary phases altered to lizardite. No talc suggesting that the sulfide is oxidized to Fe-oxides and or antigorite were detected by XRD analysis (Agrinier hydroxides during the late stage talc alteration. et al., 1996; Alt and Shanks, 1998). High d18O values of ser- The oxygen isotope compositions of serpentine pentine (around 10&) and magnetite-serpentine mineral (d18O = 2.6& to 4.4&) from Site 1268 indicate high tem- pairs indicate serpentinization at temperatures less than perature hydrothermal alteration (250–350 °C; Alt et al., 200 °C at high seawater–rock ratios (Agrinier et al., 2007)). High sulfide sulfur contents (up to 2 wt.%) and high 1996). Sulfide mineralogy is characterized by high-sulfur d34S values (4.4& to 10.8&), as well as silica metasomatism mineral assemblage (pyrite and valleriite), indicating ser- and talc alteration, suggest fluids are derived from high- pentinization occurred under low temperature (20–200 °C) temperature (>350 °C) reaction of seawater with gabbro and low H2 fugacity (Alt and Shanks, 1998). Holes 897C at depth. This contrasts to other nearby sites (1272 and and 897D are located 100 m apart, at a water depth of 1274) from ODP Leg 209, where alteration is characterized 5315 m. Hole 897C penetrated 649 m of sediment and by low temperature (<150 °C) serpentinization (e.g., high S. Ono et al. / Geochimica et Cosmochimica Acta 87 (2012) 323–340 327 d18O values up to 8.1&, and low d34S values down to The same notation is used for equilibrium fractionation 32.1&)(Alt et al., 2007). and for kinetic (biological) fractionation. The isotope enrichment factor (e) is defined as: 2.3. Seawater sulfate and unaltered MORB xe ¼ 1 xa: ð8Þ In addition to sulfur in altered oceanic basement rock Again, the common multiplication factor of 1000 is samples, sulfur from seawater and from mid-ocean ridge omitted. basalt (MORB) was extracted in order to establish the end-member compositions for the reaction models. 3.2. Sulfur isotope analysis Sulfate sulfur was extracted from six seawater samples: two samples were collected from a station near Bermuda For the altered basalt samples from the Juan de Fuca (EN408 Station 1, 1 m and 1931 m depth), and four from Ridge (Hole 1301B), sulfur was extracted from powdered near Hawaii (KOK MP5, at 12 m, 1250 m, 2500 m, and rock samples under N2 atmosphere using 6 N HCl for acid 3000 m depth, respectively) (Table 2). Sulfur was also ex- volatile sulfur (AVS) that represents acid soluble metal tracted from an unaltered seafloor basalt sample using Kiba monosulfides (e.g., FeS, NiS). Subsequent chromium reduc- extraction (Sasaki et al., 1979), a method that extracts total tion extracted the chromium reducible sulfur fraction sulfur from crystalline and glassy basalt samples using (CRS) that represents pyrite sulfur. The evolved H2S gas dehydrated tin-phosphoric acid. The sample used for this was precipitated as Ag2S. Detailed protocol for sequential was a fragment of basalt collected during R/V Atlantis extraction is described elsewhere (e.g., Alt and Shanks, Cruise AT-11-20 in 2004, at the East Pacific Rise, Segment 9°N (EPR 9°N) (sample number ALV4055-B6) (Table 3). The rock has no visual feature suggestive of alteration, Table 1 and powdered samples were washed with 0.1 N HCl to re- Results of replicate analyses of IAEA reference materials. move potential surface contamination. Sample*1 d34S(&)*2 D33S(&)*2 D36S(&)*2 S-1 (28) 0.30 (26) 0.100 (14) 0.57 (19) 3. METHOD S-2 (8) 22.24 (27) 0.043 (10) 0.16 (09) S-3 (4) 32.58 (20) 0.083 (10) 0.63 (10) 3.1. Notation Italic letters for S-1 indicate that they are defined values. *1 numbers in parentheses are the number of replicate analysis. Sulfur isotope ratios are reported using the conventional *2 numbers in parentheses are 2r standard deviations of the last delta notation: two digits. x x x d S ¼ Rsample= RVCDT 1 ð4Þ x x x 32 where Rsample and RVCDT are the isotope ratios ( S/ S, Table 2 where x = 33, 34 or 36) of the sample and VCDT (Vien- Multiple sulfur isotope compositions for seawater samples. na-Canõn Diablo Troilite reference scale), respectively. Sample d34S(&) D33S(&) D36S(&) The multiplication factor of 1000 commonly used in other KOK 1250 m 21.40 0.051 0.33 studies is omitted here because technically it belongs to KOK 2500 m 21.43 0.044 0.30 the & symbol rather than to the d notation (Coplen, KOK 3000 m 21.35 0.052 0.32 2011). This simplifies other definitions and equations in this KOK 12 m 21.26 0.053 0.37 paper as was done in Farquhar et al. (1989). The isotope ra- EN408 surface 21.31 0.048 0.33 x tios of VCDT ( RVCDT) are defined by the international ref- EN408 2000 m 21.29 0.052 0.32 erence material (IAEA-S1) to be 0.055&, 0.300&, and Average (2r)*1 21.34 (13) 0.050 (03) 0.33 (04) 33 34 36 1.14& for d S, d S and d S, respectively (Ono et al., *1 numbers in parentheses are 2r standard deviations of the last 2007). two digits. Following definitions for DxS are used (e.g., Young et al., 2002; Angert et al., 2004; Ono et al., 2006): Table 3 D33S ¼ lnðd33S þ 1Þ0:515 lnðd34S þ 1Þð5Þ Multiple sulfur isotope compositions for mid oceanic ridge basalt (ALV4055-B6). D36S ¼ lnðd36S þ 1Þ1:90 lnðd34S þ 1Þð6Þ Sample size Yield d34S(&) D33S(&) D36S(&) We choose these definitions as they allow isotope frac- (ppm) tionation to be expressed linearly (Hulston and Thode, 0.5 g 640 0.06 0.011 0.01 1965; Kaiser et al., 2004; Ono et al., 2006). However, it x 1.0 g 450 0.13 0.002 0.07 should be noted that the D S values do not behave linearly 2.0 g 310 0.40 0.010 0.01 upon mixing between sulfur from different reservoirs. 1.0 g 570 0.07 0.005 0.06 The isotope fractionation factor a between sulfide (H2S) Average 0.16 (32) 0.001 (17) 0.033 (08) *1 and sulfate (SO4) is defined as: (2r) *1 xa ¼ xR =xR ð7Þ numbers in parentheses are 2r standard deviations of the last H2S SO4 two digits. 328 S. Ono et al. / Geochimica et Cosmochimica Acta 87 (2012) 323–340

Table 4 Multiple sulfur isotope compositions for altered basalt and glass samples from Site 1301B. Extraction Depth (mbsf) S (ppm) d34S(&) D33S(&) D36S(&) Bulk rock extacted sulfur 1301B 1-1 CRS 352 452 5.1 0.005 0.09 1301B 2-2 CRS 359 410 1.8 0.004 0.09 1301B 4R2-122B CRS 369 61 0.1 0.013 0.11 1301B 4R2-122G CRS 369 45 0.1 0.013 0.13 1301B 4-2 CRS 369 365 1.0 0.004 0.12 1301B 5-2 CRS 379 451 0.2 0.017 0.08 1301B 5-2 AVS 379 169 2.1 0.008 0.04 1301B 6-2, G CRS 388 471 2.7 0.016 0.14 1301B 12R-1-28G CRS 429 476 0.8 0.011 0.17 1301B 14R1-65P CRS 435 3788 1.7 0.000 0.10 1301B 14R1-65P AVS 435 449 4.1 0.001 0.10 1301B 14R1-65B CRS 435 339 10.0 0.010 0.01 1301B 14R1-65G CRS 435 143 1.0 0.008 0.08 1301B 14-1,G CRS 435 743 5.7 0.042 0.34 1301B 14-1,B AVS 435 195 9.3 0.053 0.34 1301B 14-1,B CRS 435 35 13.7 0.057 0.46 1301B 17-16,G AVS 461 355 2.7 0.005 0.10 1301B 19R1-41G CRS 477 797 2.9 0.015 0.22 1301B 19R1-41G AVS 477 158 3.5 0.005 0.13 1301B 19R1-41B CRS 477 57 4.4 0.014 0.26 1301B 19R1-41B AVS 477 50 0.5 0.014 0.17 1301B 20-1 AVS 480 91 0.9 0.013 0.08 1301B 23R2-66P CRS 502 259 9.6 0.004 0.09 1301B 23R2-66G CRS 502 92 5.1 0.009 0.10 1301B 23R2-66G AVS 502 73 0.7 0.005 0.11 1301B 23R2-66B CRS 502 45 0.9 0.009 0.08 1301B 23-2,B AVS 502 227 0.0 0.002 0.03 1301B 23-2,G AVS 502 218 0.0 0.003 0.10 1301B 23-2,G CRS 502 122 2.1 0.013 0.23 1301B 25R1-123P CRS 511 215 3.4 0.017 0.17 1301B 25R1-123G CRS 511 94 5.2 0.014 0.12 1301B 25R1-123B CRS 511 59 1.5 0.001 0.04 1301B 26-1,G AVS 515 549 1.2 0.002 0.14 1301B 26-1,G CRS 515 666 0.9 0.005 0.14 1301B 32R3-50G CRS 553 325 1.4 0.005 0.10 1301B 32R3-50G AVS 553 268 1.6 0.008 0.12 1301B 32R3-50B CRS 553 69 4.2 0.008 0.03 1301B 36R1-53G CRS 574 178 0.1 0.003 0.18 1301B 36R1-53B CRS 574 76 1.7 0.002 0.13 1301B 36R1-53G AVS 574 72 1.5 0.004 0.12 Pyrite separates 1301B 5R-3 47-51 Py 380 n.a. 5.0 0.027 0.12 1301B 12R1-47 Py 430 n.a. 9.2 0.038 0.27 1301B 12R-1 28-34 Py 430 n.a. 12.0 0.035 0.14 1301B 12R-1 66-70 Py 430 n.a. 7.6 0.009 0.02 1301B 15R2 102-105 Py-vug 450 n.a. 5.8 0.004 0.04 1301B 18R2 71-77 Py 472 n.a. 5.7 0.002 0.01 1301B 18R-3 88-98 Py 472 n.a. 1.7 0.011 0.08 1301B 18R-3 118-123 Py-vein 472 n.a. 1.8 0.004 0.05 1301B 18R 02W 84-87 Py-ves 472 n.a. 6.2 0.012 0.21 1301B 21R1-77 Py-vug 491 n.a. 28.7 0.030 0.33 1301B 24R2-0 Py 508 n.a. 9.5 0.005 0.02 1301B 24R1 117-126 Py-frac 508 n.a. 11.3 0.016 0.05 BH = blackhalo; G = grey host rock; B = black halo; P = pyrite front. CRS, chromium reducible sulfur; AVS, acid volatile sulfur.

1998). Additionally, 12 pyrite separates from 1301B were Deep, Site 895; Iberian Margin, Site 897; Mid Atlantic used in this study, and processed using chromium reduc- Ridge 15°N. Site 1268) had been prepared and analyzed tion. Sulﬁde samples from three serpentinization sites (Hess in previous studies (Alt and Shanks, 1998; Alt et al., S. Ono et al. / Geochimica et Cosmochimica Acta 87 (2012) 323–340 329

Table 5 Multiple sulfur isotope compositions for samples from peridotite, Sites 895, 897, and 1268. Typea Depth (mbsf) S (ppm) d34S(&) D33S(&) D36S(&) Hess Deep, site 895 895E 1R3 78-81 Dunite 3.6 70 0.3 0.004 0.00 895D 3R1 95-102 Harz 27.0 168 8.1 0.013 0.04 895C 4R2 11-13 Troct 29.1 165 0.4 0.001 0.05 895D 7R2 55-57 Dunite 66.6 112 1.3 0.019 0.24 895E 7R1 0-3 Dunite 68.2 120 17.3 0.109 0.84 895E 8R1 65-69 Dunite 78.6 90 23.9 0.127 0.93 Iberian Margin, site 897 637A 27R2 100-114 Harz 249.5 90 37.2 0.011 0.18 637A 27R4 108-111 Harz 252.1 290 44.2 0.051 0.49 897C 63R2 81-85 Harz 650.5 2500 27.0 0.089 0.76 897C 64R4 0-5 Dunite 661.4 2310 14.1 0.083 0.56 897C 65R2 62-65 Harz 670.1 1240 15.1 0.080 0.59 897C 66R3 18-23 Dunite 679.8 410 20.9 0.067 0.47 897C 66R4 0-3 Gabbro 680.8 410 19.9 0.072 0.51 897C 67R2 32-38 Harz 688.6 97 30.1 0.106 0.81 897C 67R3 69-77 Gabbro 690.4 272 24.1 0.066 0.50 897C 69R1 96-100 Harz 707.2 3130 21.2 0.132 0.98 897C 70R1 54-58 Dunite 710.7 2470 21.7 0.141 1.03 897C 71R1 44-50 Harz 716.3 1320 16.6 0.135 1.02 897C 72R1 72-77 Harz 726.3 900 12.9 0.162 1.24 897C 73R3 11-14 Harz 737.2 610 6.9 0.138 0.99 897C 73R3 132-135 Dunite 738.4 1000 5.0 0.151 1.04 897D 16R1 22-27 Harz 742.2 1880 25.5 0.001 0.08 897D 16R4 84-89 Harz 746.0 290 31.9 0.068 0.51 897D 17R6 71-75 Breccia 758.5 600 15.8 0.027 0.18 897D 18R1 31-35 Harz 761.5 116 25.9 0.053 0.51 897D 19R2 105-109 Harz 773.2 1626 19.7 0.084 0.63 897D 19R5 100-105 Troct 775.7 453 17.4 0.072 0.51 897D 19R4 136-140 Harz 776.9 962 15.3 0.060 0.40 897D 21R3 57-63 Harz 792.7 1570 21.4 0.124 0.91 897D 23R2 137-140 Harz 811.7 956 7.3 0.134 0.89 897D 25R3 112-118 Harz 831.4 540 14.7 0.106 0.68 15°N Fracture Zone Mid Atlantic Ridge, site 1268 1268A 4R-1 44-55 CRS Harz 25.2 7289 8.6 0.015 0.01 1268A 18R-3 100-110 CRS Harz 95.6 669 6.0 0.037 0.19 1268A 18R-3 100-110 AVS Harz 95.6 465 6.4 0.013 0.04 1268A 19R-1 34-43 AVS Harz 97.3 854 8.6 0.012 0.09 1268A 19R-1 34-43 CRS Harz 97.3 635 10.0 0.023 0.09 1268A 20R-1 100-110 AVS Harz 102.6 740 10.1 0.007 0.04 a Harz, harzburgite; Troct, troctolite.

˚ 2007). Subsets of these Ag2S samples, mostly CRS but some column packed with molesieve 5 A followed by a column AVS, were analyzed in this study for multiple sulfur iso- packed with Hayesep Q. Isotope ratios of the purified SF6 topes. To obtain the sulfur isotope composition of seawater were analyzed by an isotope ratio mass spectrometer (Ther- sulfate, BaSO4 was precipitated from 1.5 mL of seawater by mo-electron MAT 253) by measuring the ion beams of 32 þ 33 þ 34 þ 36 þ addition of BaCl2. The BaSO4 was dried and reduced to SF5 , SF5 , SF5 , and SF5 . Replicate analyses H2S using Kiba reagent (Sasaki et al., 1979). Kiba reagent (n = 28) of a reference material, IAEA-S-1, yield 2r stan- was also used to extract sulfur from basaltic glass samples. dard deviations of 0.26&, 0.014& and 0.19& for d34S, 33 36 For all sulfur extraction procedures, H2S was precipitated D S and D S, respectively (Table 1). Similar level of 33 as Ag2S for subsequent fluorination and isotope ratio precision for D S of ca. 0.01& was achieved by previous measurements. studies at Geophysical Laboratory and University of Sulfur isotope ratios were measured by a technique sim- Maryland (Ono et al., 2006). The same level of precision 17 ilar to that described in Ono et al. (2006). Approximately is also routinely achieved for D OofO2 (e.g., Luz et al., 2mg of Ag2S were reacted under elemental fluorine 1999) and D47 value for clumped CO2 isotopologue analyses (50 torr) for over 6 h at 300 °C. The product SF6 was (e.g., Huntington et al., 2009). These errors are much purified by a gas chromatograph (GC) equipped with a smaller than those achieved in conventional single isotope 330 S. Ono et al. / Geochimica et Cosmochimica Acta 87 (2012) 323–340

Fig. 2. Histograms for d34S values of sulfide minerals measured in this study, compared with the typical ranges of d34S values for various sulfur reservoirs. Sources of d34S values are Sakai et al., (1982, 1984) for MORB and OIB, Shanks (2001) for hydrothermal sulfide (sediment- free deposits), and Sim et al. (2011a) for marine sedimentary sulfide. *Data for Western Pacific site (801C) are from Rouxel et al. (2008).

0.20 A 0.4 B 0.15

0 0.10 Peridotite

S (‰) 0.05 -0.4 S (‰) 33 36 Δ 0.00 Δ -0.8 Basalt -0.05 -1.2 -0.10 -50 -40 -30 -20 -10 0 10 20 -0.1 -0.05 0 0.05 0.1 0.15 0.2 δ 34S (‰) Δ 33S (‰) Legend Basalt sites: Juan De Fuca, 1301B (bulk, pyrite); Western Pacific*, 801C Peridotitesites: Hess Deep, 895; Iberian Margin, 897; MAR 15°N, 1268 Seawater sulfate:

Fig. 3. Plots for d34S versus D33S values (A), and D33S vs. D36S values (B) for secondary sulfides in oceanic basement rocks measured in this study. The line in (B) is D36S=7.3 D33S. Error bars represent standard deviation (1r) for eight mass spectrometer measurements of sample * SF6. Data for Western Pacific site (801C) are from Rouxel et al. (2008). ratio analysis, where they are typically about ±0.2&.Thisis handling). The strong correlations between d33S and d34S because most analytical errors for the isotope ratio analysis errors cancel out during the calculation of D33S value are due to mass-dependent processes (e.g., fluorination, gas (Ono et al., 2006). S. Ono et al. / Geochimica et Cosmochimica Acta 87 (2012) 323–340 331

4. RESULTS 500°C Equilibrium 0.516 200 100 50 4.1. Multiple sulfur isotope ratios of seawater sulfate and 20 0°C MORB 0.514 0.003

The average of the six seawater sulfate analyses yield 0.512 21.3 ± 0.1&, 0.050 ± 0.003& and 0.33 ± 0.04& (2r) θ = 0.009 34 33 36 β for d S, D S and D S, respectively, demonstrating that 0.510 seawater sulfate is isotopically homogeneous, not only in d34S(Rees et al., 1978) but also in D33S and D36S values 0.508 (Table 2). Four replicate analysis of the MORB sample from EPR 9°N yield averages of 0.16 ± 0.3, 0.001 ± 0.017, and 0.03 ± 0.08& for d34S, D33S and D36S, 0.506 0 20 40 60 80 respectively (Table 3). These results show that the D33S and D36S values of MORB are most likely identical to those 34ε (‰) of CDT within analytical precision, as assumed previously 34 (e.g., Johnston et al., 2005; Ono et al., 2007; Rouxel Fig. 4. Sulfur isotope enrichment factor ( e) and triple isotope coefficient (h) for microbial sulfate reduction (open and filled et al., 2008; Peters et al., 2010). circles), and thermodynamic equilibrium between sulfate and sulfide (solid line). Data are from Sim et al. (2011a,b) (filled 4.2. S-33 in basalt basement, Site 1301 (Juan de Fuca Ridge) circles) and Johnston et al. (2007) (open circles). Numbers on the equilibrium fractionation line represent the temperature of equi- Acid volatile sulfur (AVS) and chromium reducible sulfur librium. The grey field is a solution field for the multiple-S isotope (CRS) from bulk rock samples from Site 1301 are character- model by Farquhar et al. (2003) assuming fractionation factors of ized by depleted d34S values ranging from 1.0& to 13.7& 0.975 and 0.953 for two reduction steps of MSR (i.e., sulfate to (Fig. 2, Fig. 3-A, Table 4). Pyrite separates yield low d34S sulfite and sulfite to sulfide, respectively). The two dashed lines values that cover a range similar to those of AVS and CRS, roughly bracketing the grey field are approximated solutions for 34 except for one sample (21R1–77) that yields 28.7&. Most biological fractionation between e and h according to Eq. (13) at temperature 20 °C for the range of slope (b = 0.003 to 0.009). The D33S values are close to VCDT (±0.01&) but three bulk anal- temperatures for laboratory experiments are 20 °C for Sim et al. 33 & yses from sample 14-1 yield negative D S values (0.04 to (2011a,b) and range from 2.1 to 36.4 °C for Johnston et al. (2007). 0.06&), and three pyrite separates (5R-3 47-51, 12R-147, The data points that plot outside the field are likely due to the low- 33 12R-1 28-34) yield positive D S values (+0.03& to +0.04&). temperature (<20 °C) of the experiments by Johnston et al. (2007).

4.3. S-33 in peridotite basement sites tion, in theory, depends on a number of factors (e.g., temperature, fractionation processes) but the accuracy of D36S Sulfides (CRS and AVS) from three serpentinized peri- 34 measurements is not sufficient to derive additional informa- dotite sites studied yield a wide range of d S values, from tion for the range of D33S values found in this study. The 44.2& to 10.1&, consistent with previous results 33 36 34 correlation between D S and D S with a slope 7.3, how- (Fig. 2). This range overlaps with d S values from basaltic ever, gives confidence in the quality of data since this corre- sites when combined with the data from the basaltic base- lation is not expected for analytical artifacts due to ment of Site 801C (Rouxel et al., 2008), except for the po- 34 contamination of one of the four sulfur isotopes (see, sitive d S values found in the peridotite at Site 1268. 32 33 Ono et al., 2006). For example, preferential loss of S, Different ranges of D S values, however, are found in peri- which is suggested by Lasaga et al. (2008) during surface dotite samples when compared to the basalt samples, where 33 absorption, would produce a slope of 1.86. This is signif- peridotite samples yield relatively high D S values icantly different from the observed relationship. (0.01& up to 0.16&) as opposed to 0.06& to 0.09& for the basalt samples (Fig. 3). The Iberian Margin serpent- 5. DISCUSSION inite have a notable D33S signal with consistently high D33S values compared to basaltic Site 1301 (Fig. 3). This clearly 5.1. Multiple sulfur isotope model for ocean basement rock demonstrates that D33S can provide additional information alteration even when d34S values are identical. A simple batch reaction model was constructed in order 4.4. S-36 versus S-33 to examine the origin of the different ranges of D33S values observed in the different study sites. First, we assume that The D36S values are correlated with D33S values follow- the analyzed sulfides represent two component mixtures ing a trendline that can be described as D36S=7.31 D33S of primary (mantle-derived) sulfide and seawater sulfate- (Fig. 3-B). Ono et al. (2006) derived a slope of 6.85 as a derived secondary sulfide. Second, we assume that the second- first order approximation of D36S/D33S ratios for mass- ary sulfides are produced by sulfate reduction, whether dependent processes. The relationship measured in this inorganic or biological, in a closed system, such as inside study, 7.31 ± 0.26 (99% confidence interval) is statistically pore spaces of basement rocks. In reality, sulfate is supplied different from this first order approximation. The correla- to the basement rock by advection of basement fluids. 332 S. Ono et al. / Geochimica et Cosmochimica Acta 87 (2012) 323–340

120 0.4 A BC0.2 0.1 100 0.12 0.3 0.3 80 0.5 60 0.2 0.08 (‰) 0.7

SO Σ (‰)

Σ 4 S (‰)

SO Σ S 4 40 33 S 34

Δ 0.9 33 δ 0.1 2- 0.04 Seawater 20 Δ ΣS

0 2- ΣS 0 f = 0.99 0 Mantle -20 -0.1 1 0.8 0.6 0.4 0.2 0 1 0.8 0.6 0.4 0.2 0 -20 -10 0 10 20 30 f δ34 f SΣ (‰)

34 33 2 Fig. 5. Plots showing a typical model result for d S and D S values of total sulfide (RS ) and sulfate (SO4). (A) and (B) show the evolution of d34S and D33S, respectively, as a function of f. (C) shows d34S versus D33S values for total sulfide (RS2), and italic numbers in the diagram represent the value of f. The model was run for an initial sulfide content (mP) of 100 ppm, w/r ratio of 10, and assumed equilibrium isotope fractionations at 150 °C(a and h values of 0.964, and 0.515, respectively).

While this cannot be included in a simple batch reactor sys- 150 and 500 °C (e.g., Farquhar et al., 2003; Ono et al., tem, it is reasonable to assume for our model calculations 2007). This temperature dependence has very little effect that the isotopic composition of the advecting fluids follows on the overall results (Fig. 4). the closed system equation along the flow path. The other end-member fractionation factors are those 34 During closed system sulfate reduction, sulfur isotope associated with microbial sulfate reduction ( amsr and x ratios of product sulfide (d SSR) can be derived as follows hmsr). It has been shown that sulfate reducers can produce 34 34 x a wide range of e (=1 a ) values, from a few & 1 f a msr msr dxS ¼ðdxS þ 1Þ 1 ð9Þ to as large as 66& (e.g., Kaplan and Rittenberg, 1964; Can- SR SW 1 f field, 2001; Sim et al., 2011a,b). Various theoretical models x where d SSW is the isotope compositions of seawater sul- have been presented to explain the range of fractionation as fate, xa (x = 33, 34, or 36) is the fractionation factor be- well as the maximum isotope effect attained by microbial tween sulfate and sulfide during sulfate reduction, and f is sulfate reduction (Rees, 1973; Brunner and Bernasconi, the fraction of sulfate remaining, which is the ratio of 2005; Johnston et al., 2007). Recent experimental and theo- 32 2 32 2 3 remaining SO4 with respect to the initial SO4 . The retical studies suggest that the maximum isotope effect is model also assumes no significant isotope effect (less than determined by the equilibrium fractionation between sul- 34 a few &) during precipitation of secondary sulfide minerals fate and sulfide (i.e., aeq)(Brunner and Bernasconi, from H2S (e.g., Ohmoto and Goldhaber, 1997) or oxidation 2005; Canfield et al., 2010; Sim et al., 2011a). of pyrite (e.g., Lefticariu et al., 2006). Eq. (9) is solved for It has been shown theoretically and experimentally that 34 34 d SSR as a function of possible ranges for a and f. The hmsr differs from heq due to kinetic complications that orig- 33 same equation can also be applied to d SSR, by relating inate from isotope mass-balance and fractionation (Nor- 33a to 34a by the triple fractionation coefficient (h): throp, 1975; Farquhar et al., 2003; Johnston et al., 2005; Ono et al., 2006)(Fig. 4). In this study, the h is approx- 33a 34ah 10 msr ¼ ð Þ imated by a function: Two end-member values for 34a and h are used. One h ¼ 0:515 bð1 34e =34e Þð13Þ end-member is the equilibrium fractionation factor between msr msr eq 2- 34 34 SO4 and H2S( aeq and heq), estimated by applying con- where eeq is the isotope enrichment factor, which is a ventional isotope fractionation theory (e.g., Farquhar function of temperature (Equation (8) and (11)), b is the 34 et al., 2003; Ono et al., 2007; Otake et al., 2008). We used slope defining the correlation between h and emsr values. the formula by Ono et al. (2007): The b value ranging from 0.003 to 0.009 accommodates most experimental results as well as that expected from the- lnð34a Þ¼3424T 2 11:32T 1 þ 0:008389; and ð11Þ eq oretical models (Fig. 4). Since the maximum isotope effect 2 1 heq ¼9:460T 0:3117T þ 0:5159 ð12Þ attained by MSR at given temperature is due to equilibrium fractionation, 34e > 34e . Equation (13) is formulated in where T is the temperature in Kelvin. The estimated 34a msr eq eq such a way that it yields a kinetic end-member h value of value is consistent with experimental values between 200 msr 0.509 ± 0.003, when fractionation is minimum (i.e., and 400 °C but carries an uncertainty of up to five & at 34e 0), whereas it yields an equilibrium end-member low temperatures (<150 °C) depending upon the detailed msr of 0.515 when the maximum isotope effect at a given tem- method of calculation (e.g., Otake et al., 2008). The esti- perature is attained (i.e., 34e 34e ). Eq. (13) is an mated h value is only a weak function of temperature, msr eq eq approximate solution of a more complicated formula to de- ranging from 0.5151 to 0.5154 for temperatures between scribe a multiple step process involved in microbial sulfate reduction (e.g., Northrop, 1975; Farquhar et al., 2003; 3 The equation provides an exact solution when the ratios of Sim et al., 2011b). In the following analysis, we use the b 32 2- 32 2- value of 0.007, based on a linear regression of all available SO4 (remaining)/ SO4 (initial) are used instead of the total of all four isotopologues of sulfate. data in Fig. 4. The error due to the range of b values S. Ono et al. / Geochimica et Cosmochimica Acta 87 (2012) 323–340 333

0.25 0.25 A) Hydrothermal B) MSR 65 °C 0.1 0.2 0.2 0.3

0.15 0.15 0.1 0.3 0.6 0.1 0.6 0.1 27 13

S (‰) 40 S (‰)

33 0.05 0.9 0.05

33 0.9 Δ

Δ 0 0 f = 0.99 0 f = 0.99 ε=54 -0.05 -0.05

-0.1 -0.1 -40 -20 0 20 -40 -20 0 20 34 δ S (‰) δ 34S (‰)

0.25 0.25 0.3 D) MSR 2 °C 0.3 0.1 C) MSR 40 °C 0.1 0.6 0.2 46 0.2 ε=61 ε=75 56 37 0.15 0.6 30 0.15 0.9

0.1 15 0.1 19 S (‰) S (‰) 0.05 0.05 33 33 0.9 0 Δ Δ 0 0 0

-0.05 f = 0.99 -0.05 f = 0.99 -0.1 -0.1 -40 -20 0 20 -40 -20 0 20 δ 34S (‰) δ 34S (‰)

Fig. 6. Model results in d34S versus D33S coordinate system. (A) represents solutions for hydrothermal sulfides produced by thermochemical sulfate reduction. The contours without labels represent equilibrium fractionation at 150, 200, 250, 400 and 1300 °C from upper left to lower right, and the contours with italic numbers show various f values. (B), (C), and (D) are solutions for microbial sulfate reduction at 65, 40 and 2 °C, respectively. The contours represent different values for e and f. Italic numbers are for the value of f. Other input parameters are the initial sulfide content (mP) of 100 ppm, w/r ratio of 10. Symbols for the data are identical to those in Fig. 3.

34 34 (0.006 ± 0.003) is only critical when emsr is small, and the respect to VCDT. The actual d S values for MORB and area of the solution field is not sensitive to the b value. Sen- OIB may vary by ±2& (Fig. 2) but this does not affect sitivity of different b values to the model solution is shown the overall conclusions. in Supplementary Fig. A1. Eqs. (9) and (14) are solved for d34S and d33S, and solved 33 34 Based on the first assumption of the model, the sulfide for D S using Eq. (5), as a function a, f, h, mp, and (w/r). minerals extracted from altered basalt and peridotite rocks These are five variables, but mP and (w/r) are related represent the mixing of primary sulfide of mantle origin (according to Eq. (14) and (15)) and can be combined to with secondary sulfide produced either by microbial or by give one dimensionless quantity (e.g., (w/r) mSW/mP), and thermochemical sulfate reduction. The isotopic composi- h is related to 34a according to Eq. (13). We can then solve x 34 tion of the total sulfide (d SR) can be solved by simple iso- the model with a and f as main variables, and carry out tope mass-balance: sensitivity tests for various (w/r) ratios (Fig. 6 and m Fig. A1 in Spplementary material). dxS ¼ SR dxS ð14Þ It should be noted that we do not specifically consider R m þ m SR SR P potential effects of sulfur disproportionation. Sulfur dispro- where mP is the concentration primordial sulfide (i.e., sul- portionation reactions, either abiotic (Andrews, 1979)or fide sulfur originated from the mantle), and mSR is the con- biotic (e.g., Canfield and Thamdrup, 1994; Habicht and centration of secondary sulfide from sulfate reduction. mSR Canfield, 2001; Johnston et al., 2005; Riedinger et al., is given by: 2010), would proceed via multiple sulfur reservoirs of inter- mediate redox levels (e.g., S O2,SO2). The exact solu- m 1 f m w=r 15 2 4 3 4 SR ¼ð Þ SWð ÞðÞ tion is difficult to constrain without a thorough where mSW is the sulfate concentration in seawater understanding of detailed chemical pathways and isotope (900 ppm), and w/r represents the water to rock ratio by effects. Rapid isotope exchange between H2S and polysul- 2 weight. Equation (14) assumes that the isotopic composi- fide (Sn ) (e.g., Fossing and Jørgensen, 1990) also compli- tion of primary sulfur (i.e., mantle sulfur) is 0& with cates a model solution. We note here that there are some 334 S. Ono et al. / Geochimica et Cosmochimica Acta 87 (2012) 323–340

30 indications that sulfur disproportionation may not be re- A) Hydrothermal 0.9 0.6 0.3 quired to explain large (>47&) sulfur isotope effects, and 20 1300 large sulfur isotope effects may be produced by sulfate f = 0.99 reduction only (Canfield et al., 2010; Sim et al. 2011a). 10 Johnston et al., (2005) has reported an interesting observa- 400 tion that microbial sulfur disproportionation produce a un- 0 250 ique range of h values higher than 0.515 between sulfate 200 and sulfide (i.e., two products). However, the h values are (‰) -10 T=150 Σ

S smaller than 0.515 when compared between reactant and 34 0 δ -20 product (e.g., S and H2S). This suggests that the low D33S values measured for Site 1301, for example, may re- -30 quire more than a simple one-step sulfur disproportionation process because it would produce sulﬁdes that have -40 higher D33S values compared to starting S0 (or other inter- -50 mediate sulfur compounds). Therefore, in this study, we 10 100 1000 10000 will test the scenario for single step microbial sulfate reduction. A full model including sulfur disproportionation may mΣ (ppm) be constructed in future studies when more experimental 30 data become available. B) MSR 40 °C 0.9 0.6 0.3 20 0 5.1.1. Model solution, characteristics and sensitivity f = 0.99 Fig. 5 shows the results of calculations for the evolution 10 34 of isotope compositions for total bulk rock sulﬁde (d SR 15 33 and D SR) during sulfate reduction. As the sulfate reduc- 0 34 tion proceeds (i.e., f decreases), the d SR value decreases

(‰) 30 34 Σ initially from the mantle value (d S = 0), due to the addi-

S -10 34 34 tion of S-depleted secondary sulfide formed by sulfate δ -20 reduction (this applies if 34e >21&, i.e., d34S value of sea- 46 34 water sulfate). As sulfate reduction progresses, d SR in- -30 34 creases because of the increase in d SSR due to closed ε=61 33 system effects (Fig. 5-A). The evolution of the D SR value -40 33 is a complex function of f but relatively high D SR values -50 are expected for low f values (Fig. 5-B). Fig. 5-C demon- 10 100 1000 10000 strates the most critical aspect of the model result, which is that there is a range of d34S values where there are two mΣ (ppm) solutions for the value of f at a given d34S value. For exam- 34 Fig. 7. Model results are shown in a sulfur content (mR) versus ple, a d SR value of 8& can be the result of the addition 34 d SR coordinate system. (A) shows the hydrothermal solution of isotopically depleted sulfide produced at either f 0.99 33 (corresponding to Fig. 6-A), and (B) shows solution for microbial or f 0.7. The latter sulfide, however, yields higher D SR sulfate reduction at 40 °C (corresponding to Fig. 6-C). Symbols for value so that it can be distinguished from the former exam- the data are identical to those in Fig. 3. ple. This illustrates the usefulness of analyzing multiple

0.1 0.1 A) Hydrothermal 0.3 B) MSR 65°C 27 40 13 T=150 200 250 0.6 0.9 0.05 0.9 0.05 400 f = 0.99 ε=54 S (‰) 0 1300 S (‰) 0 0 33 33 f = 0.99 Δ 21R-1-77 Δ

-0.05 -0.05

ε=75 14-1 -0.1 -0.1 -30 -20 -10 0 10 -30 -20 -10 0 10 δ 34S (‰) δ 34S (‰)

Fig. 8. The model results and data for basaltic site 1301. The diagrams are identical to Fig. 6-A and B but in diﬀerent axis ranges, and data are shown exclusively for 1301. Four samples, in particular, (three from 14-1 and one from 21R-1-77) plot outside of the MSR ﬁeld at 65 °C. The dashed line represents the maximum enrichment factor for MSR at 2 °C(e =75&). S. Ono et al. / Geochimica et Cosmochimica Acta 87 (2012) 323–340 335

33 34 sulfur isotopes and shows the unique aspect of D SR val- sponds to the solution for Fig. 6-C. The mR versus d SR ues, which provide information that cannot be attained diagrams provide a qualitative assessment of the sulfur iso- 34 by the analysis of d SR only. tope systematics but the solution field depends critically on x Fig. 6 shows the sensitivity of the model to a and f, the initial (i.e., primary) sulfide content (mP). This high- 34 33 assuming mP of 100 ppm and (w/r) of 10 (i.e., (w/r) mSW/ lights the advantage of the new d S-D S approach, which mP = 90). These values are chosen because they yield total is much less sensitive to the initial sulfide content. sulfide contents (mR) that are consistent with measurements, as discussed later (Fig. 7). The initial magmatic sul- 5.2. S-33 constraints on the alteration of oceanic basement fur content of basaltic lava at Site 1301 is estimated to be rocks 1390 ± 80 ppm based on the analysis of six basalt glass samples (Supplementary material). Previous studies for 5.2.1. Basalt site other basaltic basement sites showed that the loss of mag- First, we examine if bulk sulfide extracts or pyrite iso- matic sulfur during degassing and oxidation followed by la- lates from basalt Site 1301 are consistent with precipitation ter sulfur addition by precipitation of secondary sulfide from a high-temperature hydrothermal fluid during the minerals (Rouxel et al., 2008; Alt and Shanks, 2011). The early stages of basement rock alteration. The model results apparently small contribution of magmatic sulfur is also show that only 14 samples out of 52 plot in a hydrothermal due to sulfur extraction technique (AVS and CRS) that field (Fig. 8-A). That is, the sulfur isotope ratios (d34S and preferentially extracts sulfur in sulfide minerals, as opposed D33S) of these 14 samples can be explained by the addition to sulfur dissolved in basaltic glass. As shown later, the of secondary sulfide formed by sulfate reduction with a 34e 34 33 34 choice of mp is not critically sensitive to the d S-D S mod- value less than 37& (i.e., eeq at 150 °C). The rest of sam- el solution unless a very low w/r ratio is assumed. ples require sulfate reduction at temperature lower than Fig. 6-A shows the model solutions for thermochemical 150 °C, which suggests microbially mediated sulfate sulfate reduction, in which aeq are calculated for tempera- reduction. tures from 150 °C to 1300 °C. The upper temperature limit The question arises whether these potential microbial was tested to model the case of no isotope fractionation sulfide components are the result of in-situ MSR within 34 (i.e., aeq = 1). The solution for different f and e (=1 a) the oceanic basement or these S-depleted sulfides were values are shown as contours. The model results for micro- produced by MSR in overlying sediments and introduced bial sulfate reduction (MSR) are shown in Fig. 6-B, C, and into basement rocks as hydrogen sulfide. The current tem- D, for temperatures of 65, 40 and 2 °C. As discussed later, perature of the basement fluid is 64.5 °C (measured at Site the temperatures of 65 and 40 °C represent those of current 1026 which is about 2 km away from Site 1301; Davis and basement fluids at Sites 1301 and 897, respectively (Sawyer Becker, 2002). However, the basement temperature was et al., 1994; Davis and Becker, 2002), whereas 2 °C is the lower in the past as documented by the d18O values of vein temperature of bottom seawater at Site 1301 (Davis and carbonate, which suggests an alteration fluid temperature Becker, 2002). The model results show that the biological of 3–50 °C(Coggon et al., 2004). The basement tempera- solution fields occupy significantly larger areas compared ture could be as low as near ambient seawater soon after to the hydrothermal fields. This is due to larger fraction- emplacement. 34 ation factors attained by the lower temperature microbial When applying eSR and hSR for MSR at 65 °C (i.e., processes. Fig. 6-C and D illustrate how biogenic sulfate current basement temperature), the expanded solution 33 reduction can produce negative D SR values, which is not fields can accommodate about a half of the analyzed sam- expected for thermochemical equilibrium fractionation. ples (29 out of 52), with a considerable number of samples The model result is most sensitive to the enrichment fac- plotting outside the field (Fig. 8-B). Extending the field for tor (34e), and the extent of sulfate reduction (f). The water/ MSR at 2 °C (dashed line in Fig. 8-B), which is the bottom rock ratio (or mP) has a secondary importance. In fact, the seawater temperature at Site 1301 (Davis and Becker, extent of the solution field is not sensitive to the w/r ratios 2002), can accommodate 39 samples. In particular, analyses unless a w/r ratio of lower than 10 is assumed (Supplemen- of three different alteration zones from one sample (1301B tary material, Fig. A2). This mathematical behavior is ex- 14-1) plot on the edge of the field defined by MSR at 34 pected from Equation (14) because when mSR mP, 2 °C. The grey host rock yields less negative d S values 34 34 d SR d SSR. Based on the oxygen isotope ratios of ser- compared to the visually altered alteration halo samples pentine minerals, Alt et al. (2007) estimated the w/r ratio (Table 4). This suggests that sulfide in these samples could to be higher than 10 for low temperature (<150 °C) serpent- be sourced from microbial activity within the basement but inization at MAR-15 °N. Much higher w/r ratios of this would have had to happen during the very early stages 780 ± 280 are estimated for basalt basement rock (Bach of basement history when the lavas were cooler. Alterna- and Edwards, 2003). tively, the secondary sulfides could be sourced from micro- In Fig. 7, the model solutions are shown in mR (total sul- bial sulfate reduction occurring in overlying sediments. This 34 fur contents) versus d SR coordinates, similar to a sche- signature appears to be limited to samples 14-1 and 21R-1- matic diagram used previously (e.g., Alt and Shanks, 77, suggesting possible localized fluid flows from the sedi- 1998). The equilibrium end-member model is shown in ment cover. Some samples (10 out of 52) plot outside the Fig. 7-A, which corresponds to the model solution for MSR-2 °C field (plot d34S=±2& and D33S < 0), but the hydrothermal sulfate reduction (Fig. 6-A). Fig. 7-B is a significance of these data is unclear since the signal is about solution for biological end-member at 40 °C, and corre- the size of the analytical error (Fig. 8-B). 336 S. Ono et al. / Geochimica et Cosmochimica Acta 87 (2012) 323–340

5.2.2. Peridotite sites metasomatism also indicate gabbro in the source region of The temperature of basement fluids was not directly hydrothermal fluids at this site (Bach et al., 2004). measured for Site 897 (Iberian Margin). Sawyer et al. (1994) report a temperature of 13.1 °C at 214.4 mbsf and 5.3. Mass balance of hydrogen production and sulfate a geothermal gradient of 43 mK/m in the sediments overly- reduction ing Site 897. Extrapolating the geothermal gradient yields 36–39 °C at the basement/sediment interface depth (745 m Sulfate reduction coupled with H2 oxidation can be writ- at Site 897C and 837 m at Site 897D). Current basement flu- ten as: ids at Site 897, therefore, are likely to be 40 °C or higher. 2 þ SO4 þ 4H2 þ 2H ! H2S þ 4H2O ð16Þ Twenty-two out of 25 samples plot in the MSR-40 °C 34 field with e values between 61& and 30&, and f between The product H2S can precipitate pyrite (FeS2) or pyr- 2+ 3+ 0.99 and 0.6 (Fig. 6), suggesting the majority of samples rhotite (Fe1xS) via reaction with Fe or Fe minerals. may carry sulfide produced by in-situ sulfate reduction. Several pyrite precipitation pathways are possible (e.g., Samples plotting outside the field include two samples from Schoonen, 2004), but we write a reaction with FeOOH Site 637 (637A 27R2 and 27R4). These two samples, how- (goethite) since the reaction with ferric iron minerals is 2+ ever, plot within the field defined by MSR at 2 °C, suggest- much faster than that with Fe silicate minerals (e.g., Can- ing that these sulfides were formed when the basement field et al., 1992). temperature was low or sourced from overlying sediments 2FeOðOHÞþ4H2S ! 2FeS2 þ 2H2O þ H2 ð17Þ at lower temperatures (Table 5; Fig. 6-C). Such low temperature alteration at Site 637 is consistent with the d18O value Combining these two reactions with the hydrolysis of 2+ of carbonate suggesting alteration temperature as low as Fe -silicate minerals (reaction 2), the overall reaction of 10 °C(Agrinier and Girardeau, 1988). One sample (897D sulfate reduction for (2), (16), and (17) can be written as: 25R3) has distinctly different values than the other samples 2 þ 15FeO þ 2SO4 þ 5H2O þ 4H for Site 897. This sample is from the deepest depth (831.4 mbsf), and shows a positive d34S value of 14.7& and a ! FeS2 þ 14FeOðOHÞð18Þ 33 D S value of 0.106&, These isotope compositions suggest This reaction stoichiometry indicates that precipitation sulfate reduction at very low f value of about 0.1 (Fig. 6- of one mole of pyrite would require oxidation of 15 mol C). This is consistent with a previous study that shows of ferrous iron. This suggests that pyrite formation would 34 the trend of d S versus depth at Site 897, from negative require pervasive oxidative weathering provided that there values over the upper 100 m of basement to positive values is no supply of external reductant (e.g., organic materials). at depths greater than 120 m, a trend thought to reflect Mid-Ocean Ridge Basalt (MORB) typically contains ca. closed system behavior at depth (Alt and Shanks, 1998). 8 wt.% Fe, of which 2.4 ± 1.3 wt.% are oxidized in ridge When compared to the basalt site at 1301, data for the peri- flanks within the initial 10–15 Ma of their formation (Bach dotite site at 897 are characterized by a range of f values and Edwards, 2003). Bach and Edwards (2003) suggested that extends to lower values (0.99–0.6 for 897 as opposed that approximately half of the Fe oxidation can potentially to >0.99 for Site 1301). The significance of these results is be coupled to H2 production. This corresponds to discussed in the next section. 210 ± 130 mmol Fe/kg basalt. According to Reaction Two of the samples from Hess Deep (Site 895 – samples (18), this translates to the sulfate reduction capacity of ba- 895E 7R1 and 895E 8R1) fall into the range defined by the salt to be 29 ± 17 mmol SO4/kg basalt (or 770 mg S/kg ba- 34 Iberian Margin data ( e =50±5&, and f = 0.9 ± 0.1) salt). Therefore, sulfur addition up to 770 ppm S is possible 34 whereas three samples plot close to 0& (d S=0.3& to through Fe2+ hydration, and does not require an external 33 1.3&; D S = 0.00& to 0.01&). The latter sulfides are reductant. This is consistent with the relatively low sulfur mostly derived from primary (mantle) sulfur. One sample enrichment measured in basalt from Site 1301 (Fig. 7, Ta- (895D 3R1) yields isotope values between the two end- ble 1), where most (45 out of 50) samples yield sulfide less members. than 500 ppm. If w/r ratios for low temperature basalt The samples from MAR-15°N (Site 1268) yield a unique basement alteration are as high as 780 (Bach and Edwards, 33 34 34 D S-d S relationship, with positive d S (6.0–10.1&) and 2003), this implies that only 0.11% of sulfate is reduced to 33 33 D S values (0.007–0.037&)(Fig. 6). These ranges of D S- sulfide (i.e., 1-f = 770 ppm /900 ppm/780 = 0.0011). This 34 d S values are similar to those measured in seafloor hydro- explains the open system sulfate reduction (f > 0.99) esti- thermal sulfide deposits (Ono et al., 2007). We interpret these mated from our D33S model (Fig. 8-A and B). 33 D S data to indicate a hydrothermal origin of the sulfides at The typical Fe content of peridotite is also 7 to 8 wt.% low w/r ratios at high temperatures (as high as 400 °C). These but as much as two thirds of the iron in olivine can be oxi- data can be explained by low w/r ratio of 0.1, and near-quan- dized to Fe3+ (reaction 1) so that the maximum oxidizable titative sulfate reduction (f 0) (Supplementary Fig. A2). Fe is 5.0 ± 0.3 wt.%. This is higher than that of basalt by a 34 Based upon the high sulfur content and high d S values of factor of five, and the corresponding reducing capacity is 18 these samples, together with relatively low d O of the serp- higher for peridotite compared to basalt. The maximum entinites from the same site, Alt et al. (2007) suggested that H2 production capacity of peridotite through serpentiniza- the sulfides were precipitated from a hydrothermal fluid that tion (by the formation of magnetite) would be 430 mmol acquired sulfur by high-temperature (>350 C) reaction of ° H2/kg peridotite. This very crude estimate is roughly con- seawater with gabbro at depth. Talc alteration and silica sistent with a more rigorous thermodynamic analysis by S. Ono et al. / Geochimica et Cosmochimica Acta 87 (2012) 323–340 337

McCollum and Bach (2009), suggesting 360 mmol H2/kg at hydrothermal sulfide. The model shows that biogenic sul- 300 °C (at w/r = 10). The H2 production capacity is ex- fides occupy a much larger field, due to the larger fraction- pected to be lower than this value during low temperature ation factor that is intrinsic to biological sulfate reduction serpentinization (100 mmol H2/kg at 100 °C) due to in- (Fig. 9). Because the maximum isotope effect attained by creased partitioning of Fe2+ into brucite and serpentine so- MSR in the model is set by the equilibrium isotope effect lid solutions (McCollum and Bach, 2009). These estimates between sulfate and sulfide, the size of the field for MSR translate into sulfate reduction capacities of 860 and is primarily a function of temperature. This provides a un- 3100 ppm S for serpentinized peridotite at 100 and ique and critical constraint for the origin of secondary sul- 300 °C, respectively. These values yield f values of 0.90 to fide in basement rocks because basement rocks are 0.66, respectively, assuming w/r ratios of 10, consistent with generally warmer than the overlying sediments due to geo- the range of f values derived from the new d34S-D33S model thermal gradient of the non-permeable sediment cover. (Fig. 6-C). This simple mass balance calculation supports the D33S data that indicate closed system sulfate reduction 6. CONCLUSION during the serpentinization reaction. We show that D33S values of sulfide extracted from al- 5.4. S-33 as a biosignature tered basement rocks add additional information to untan- gle multiple origins of sulfides in altered oceanic basement From the comparison of D33S values of biogenic sulfide rocks, namely primary (mantle-derived) sulfide, hydrother- of various ages with hydrothermal sulfides, Ono et al. mal sulfide and microbial sulfide (in-situ in the basement or 33 (2007) suggested that the biogenic sulfides carry D33S values from overlying sediments). In particular, the D S value is higher than seawater because bacterial fractionation yields sensitive to open versus closed system sulfate reduction, h values of less than 0.515 (see also Peters et al., 2010). It and can be used to estimate isotope fractionation factors was recognized in later studies, however, that the mixing intrinsic to sulfate reduction processes by correcting the ef- between highly 34S-depleted microbial pyrite and primor- fect of closed system processes. 33 dial sulfide can produce negative D33S values because mix- The D S data suggest addition of secondary sulfide by ing is expressed non-linearly in D33S values (e.g., Rouxel an open system sulfate reduction for two basalt basement et al., 2008; Shen et al., 2011). sites (Sites 1301 and 801c), contrasting with closed system The data presented in this contribution expand the mul- sulfate reduction at two peridotite sites (Site 895 and tiple sulfur isotope systematics presented in these former 897). The closed system sulfate reduction in these two sites studies. We formulate a model that can be applied to both is consistent with the high H2 production capacity of ser- 34 33 open and closed systems (Fig. 9), that is, isotope systemat- pentinization reactions. We present a d S-D S diagram ics during mixing of two sulfur pools as well as during to show the solution field for in-situ microbial sulfate reduc- closed system processes. The new model presented in this tion. By using this new proxy, we identified a set of samples study forms a defined field in d34S-D33S diagrams for hydro- from peridotitic basement rocks that carry sulfide sourced thermal sulfide, which reflects the mixing of primordial and from sulfate reduction under basement rock temperature, and conclude that these are likely formed by microbial sulfate reduction within oceanic basement rocks. These micro- 0.4 bial d34S-D33S signatures contrast with those measured for a third peridotite site (Site 1268), where sulfide minerals are 0.3 derived from hydrothermal sulfate reduction. Basaltic base- sediment biogenic 2 °C ment rocks also carry sulfide from microbial sulfate reduction but isotope data also suggest that sulfide in a few basalt 0.2 40 °C in-situ biogenic samples are formed during a very early stage of alteration S (‰) or introduced from the overlying sediments, whereas a 33

Δ 150 °C few other samples show evidence of open system thermo- 0.1 hydrothermal chemical sulfate reduction. ACKNOWLEDGEMENTS 0 This work was supported partially by NSF OCE-0753126 to S. Ono, O. Rouxel, and J. Alt, and by a Swiss National Science Foun- -0.1 -40 -20 0 20 dation postdoctoral grant to N. Keller. Authors thank Ed Boyle for providing seawater samples, and William Olszewski for assis- δ34S (‰) tance with sample analysis. Reviews by Huiming Bao, Hiroshi Fig. 9. The d34S-D33S solution fields for sulfate reduction at Ohmoto, Ed Ripley significantly improved the manuscript. various temperatures. The hydrothermal field is defined by the temperature above 150 C, microbial sulfate reduction field is ° APPENDIX A. SUPPLEMENTARY DATA defined here by the temperature between 150 and 40 °C, where most data from peridotite basements plot. Data plots within the solution field between 40 and 2 °C may indicate sulfide produced at Supplementary data associated with this article can be ambient seafloor temperature. Data symbols are the same as in found, in the online version, at http://dx.doi.org/10.1016/ Fig. 3. j.gca.2012.04.016. 338 S. Ono et al. / Geochimica et Cosmochimica Acta 87 (2012) 323–340

REFERENCES Canfield D. E. (2001) Biogeochemistry of Sulfur Isotopes. Rev. Mineral. Geochem. 43, 607–636. Agrinier P., Cornen G. and Beslier M. (1996) Mineralogical and Canfield D. E., Farquhar J. and Zerkle a L. (2010) High isotope oxygen isotopic features of serpentinites recovered from the fractionations during sulfate reduction in a low-sulfate euxinic ocean/continent transition in the Iberia Abyssal Plain. Proc. ocean analog. Geology 38, 415–418. ODP 149, 541–552. Canfield D. E., Raiswell R. and Bottrell S. (1992) The reactivity of Agrinier P. and Girardeau J. (1988) Hydrothermal alteration of the sedimentary iron minerals toward sulfide. Am. J. Sci. 292, 659- peridotites cored at the ocean/continent boundary of the 659. Iberian margin: petrologic and stable isotope evidence. Proc. Coggon R. M., Teagle D. A. H., Cooper M. J. and Vanko D. A. ODP Sci. Results 103, 225–234. (2004) Linking basement carbonate vein compositions to Alt J. C. (2004) Alteration of the upper oceanic crust at low porewater geochemistry across the eastern flank of the Juan temperatures. In Hydrogeology of the Oceanic Lithosphere (eds. de Fuca Ridge, ODP Leg 168. Earth Planet. Sci. Lett. 219, 111– E. E. Davis and H. Elderfield). Cambridge Univ Press, 128. Cambridge, pp. 495–533. Coplen T. B. (2011) Guidelines and recommended terms for Alt J. C., Anderson T. F. and Bonnell L. (1989) The geochemistry expression of stable-isotope-ratio and gas-ratio measurement of sulfur in a 1.3 km section of hydrothermally altered oceanic results. Rapid Commun. Mass Spectrom. 25, 2538–2560. crust, DSDP Hole 504B. Geochim. Cosmochim. Acta 53, 1011– Cowen J. P., Giovannoni S. J., Kenig F., Johnson H. P., Butterfield 1023. D., Rappe´ M. S., Hutnak M. and Lam P. (2003) Fluids from Alt J. C. and Shanks W. C. (1998) Sulfur in serpentinized oceanic aging ocean crust that support microbial life. Science 299, 120– peridotites: Serpentinization processes and microbial sulfate 123. reduction. J. Geophys. Res. 103, 9917–9929. Davis E. E. and Becker K. (2002) Observations of natural-state Alt J. C. and Shanks W. C. (2011) Microbial sulfate reduction and fluid pressures and temperatures in young oceanic crust and the sulfur budget for a complete section of altered oceanic inferences regarding hydrothermal circulation. Earth Planet. basalts, IODP Hole 1256D (eastern Pacific). Earth Planet. Sci. Sci. Lett. 204, 231–248. Lett. 310, 73–83. Delacour a., Fruhgreen G., Bernasconi S. and Kelley D. (2008) Alt J. C., Shanks W. C., Bach W., Paulick H., Garrido C. J. and Sulfur in peridotites and gabbros at Lost City (30 N, MAR): Beaudoin G. (2007) Hydrothermal alteration and microbial implications for hydrothermal alteration and microbial activity sulfate reduction in peridotite and gabbro exposed by detach- during serpentinization. Geochim. Cosmochim. Acta 72, 5090– ment faulting at the Mid-Atlantic Ridge, 15 200N (ODP Leg 5110. 209): a sulfur and oxygen isotope study. Geochem. Geophys. Elderfield H. and Schultz A. (1996) Mid-ocean ridge hydrothermal Geosyst. 8, Q08002. http://dx.doi.org/10.1029/2007GC001617. fluxes and the chemical composition of the ocean. Annu. Rev. Anderson R., Chapelle F. and Lovley D. (1998) Evidence against Earth Planet. Sci. 24, 191–224. hydrogen-based microbial ecosystems in basalt aquifers. Sci- Elderfield H., Wheat C., Mottl M., Monnin C. and Spiro B. (1999) ence 281, 976–977. Fluid and geochemical transport through oceanic crust: a Andrews A. (1979) On the effect of low-temperature seawater- transect across the eastern flank of the Juan de Fuca Ridge. basalt interaction on the distribution of sulfur. Earth Planet. Earth Planet. Sci. Lett. 172, 151–165. Sci. Lett. 46, 68–80. Farquhar G., Ehleringer J. and Hubick K. (1989) Carbon isotope Angert A., Cappa C. D. and DePaolo D. J. (2004) Kinetic 17O discrimination and photosynthesis. Annu. Rev. Plant Physiol. effects in the hydrologic cycle: indirect evidence and implica- Pant Mol. Biol. 40, 503–537. tions1. Geochim. Cosmochim. Acta 68, 3487–3495. Farquhar J., Johnston D. T., Wing B. A., Habicht K. S., Canfield Bach W. and Edwards K. J. (2003) Iron and sulfide oxidation D. E., Airieau S. and Thiemens M. (2003) Multiple sulphur within the basaltic ocean crust: implications for chemolitho- isotopic interpretations of biosynthetic pathways: implications autotrophic microbial biomass production. Geochim. Cosmo- for biological signatures in the sulphur isotope record. Geobi- chim. Acta 67, 3871–3887. ology 1, 27–36. Bach W., Garrido C. J., Paulick H., Harvey J. and Rosner M. Fisher A. T. (2005) Marine hydrogeology: recent accomplishments (2004) Seawater-peridotite interactions: first insights from ODP and future opportunities. Hydrol. J. 13, 69–97. Leg 209, MAR 15 N. Geochem. Geophys. Geosyst. 5. http:// Fisher A. T., Davis E., Hutnak M., Spiess V., Zu¨hlsdorff L., dx.doi.org/10.1029/2004GC000744. Cherkaoui A., Christiansen L., Edwards K., Macdonald R. and Bach W., Paulick H., Garrido C. J., Ildefonse B., Meurer W. P. and Villinger H., et al. (2003) Hydrothermal recharge and discharge Humphris S. E. (2006) Unraveling the sequence of serpentini- across 50 km guided by seamounts on a young ridge flank. zation reactions: petrography, mineral chemistry, and petro- Nature 421, 618–621. physics of serpentinites from MAR 15 N (ODP Leg 209, Site Fisher, A.T., Urabe, T., Klaus, A., and Expedition 301 Scientists, 1274). Geophys. Res. Lett. 33, 4–7. 2005. Proceedings of the Integrated Ocean Drilling Program Banerjee N. R. and Muehlenbachs K. (2003) Tuff life: bioaltera- 301: College Station, TX (Ocean Drilling Program. tion in volcaniclastic rocks from the Ontong Java Plateau. Fisk M. R., Giovannoni S. J. and Thorseth I. H. (1998) Alteration Geochem. Geophys. Geosyst. 4, 1037. http://dx.doi.org/10.1029/ of oceanic volcanic glass: textural evidence of microbial 2002GC000470. activity. Science 281, 978–980. Boillot, G., Winterer, E.L., Meyer, A.W., et al., 1987. Proceeding Furnes H. and Staudigel H. (1999) Biological mediation in ocean ODP Initial Reports, 103: College Station, TX (Ocean Drilling crust alteration: how deep is the deep biosphere? Earth Planet. Program). http://dx.doi.org/10.2973/odp.proc.ir.103.1987. Sci. Lett. 166, 97–103. Brunner B. and Bernasconi S. (2005) A revised isotope fraction- Fossing H. and Jørgensen B. B. (1990) Sulfur isotope exchange ation model for dissimilatory sulfate reduction in sulfate reactions with radiolabeled in anoxic seawater. Biogeochemistry reducing bacteria. Geochim. Cosmochim. Acta 69, 4759–4771. 9, 223–245. Canfield D. and Thamdrup B. (1994) The production of 34S- Habicht K. S. and Canfield D. E. (2001) Isotope fractionation by depleted sulfide during bacterial disproportionation of elemen- sulfate-reducing natural populations and the isotopic compo- tal sulfur. Science 266, 1973. sition of sulfide in marine sediments. Geology 29, 555. S. Ono et al. / Geochimica et Cosmochimica Acta 87 (2012) 323–340 339

Hayes J. M. and Waldbauer J. R. (2006) The carbon cycle and Ohmoto H. and Lasaga A. C. (1982) Kinetics of reactions between associated redox processes through time. Philos. Trans. R. Soc. aqueous sulfates and sulfides in hydrothermal systems. Geo- London, B 361, 931–950. chim. Cosmochim. Acta 46, 1727–1745. Hulston J. and Thode H. G. (1965) Variations in the S33,S34, and Ono S., Shanks W., Rouxel O. and Rumble D. (2007) S-33 S36 contents of meteorites and their relation to chemical and constraints on the seawater sulfate contribution in modern nuclear effects. J. Geophys. Res. 70, 3475–3484. seafloor hydrothermal vent sulfides. Geochim. Cosmochim. Acta Huntington K. W., Eiler J. M., Affek H. P., Guo W., Bonifacie M., 71, 1170–1182. Yeung L. Y., Thiagarajan N., Passey B., Tripati A., Dae¨ron M. Ono S., Wing B., Johnston D. T., Farquhar J. and Rumble D. and Came R. (2009) Methods and limitations of “clumped” (2006) Mass-dependent fractionation of quadruple stable sulfur

CO2 isotope (D47) analysis by gas-source isotope ratio mass isotope system as a new tracer of sulfur biogeochemical cycles. spectrometry. J. Mass Spectrom. 44, 1318–1329. Geochim. Cosmochim. Acta 70, 2238–2252. Johnston D. T., Farquhar J. and Canfield D. E. (2007) Sulfur Otake T., Lasaga A. C. and Ohmoto H. (2008) Ab initio isotope insights into microbial sulfate reduction: when microbes calculations for equilibrium fractionations in multiple sulfur meet models. Geochim. Cosmochim. Acta 71, 3929–3947. isotope systems. Chem. Geol. 249, 357–376. Johnston D. T., Wing B. A., Farquhar J., Kaufman A. J., Strauss Orcutt B. N., Sylvan J. B., Knab N. J. and Edwards K. J. (2011) H., Lyons T. W., Kah L. C. and Canfield D. E. (2005) Active Microbial ecology of the dark ocean above, at, and below the microbial sulfur disproportionation in the mesoproterozoic. seafloor. Microbiol. Mol. Biol. Rev. 75, 361–422. Science 310, 1477–1479. Peters M., Strauss H., Farquhar J., Ockert C., Eickmann B. and Kaiser J., Ro¨ckmann T. and Brenninkmeijer C. A. M. (2004) Jost C. L. (2010) Sulfur cycling at the Mid-Atlantic Ridge: A Contribution of mass-dependent fractionation to the oxygen multiple sulfur isotope approach. Chem. Geol. 269, 180–196. isotope anomaly of atmospheric nitrous oxide. J. Geophys. Res. Rees C. (1973) A steady-state model for sulphur isotope fraction- 109, 1–11. ation in bacterial reduction processes. Geochim. Cosmochim. Kaplan I. and Rittenberg S. (1964) Microbiological fractionation Acta 37, 1141–1162. of sulphur isotopes. J. Gen. Microbiol. 34, 195–212. Rees C., Jenkins W. and Monster J. (1978) The sulphur isotopic Kelley D., Karson J. A., Fru¨h-Green G. L., Yoerger D. R., Shank composition of seawater sulphate. Geochim. Cosmochim. Acta T. M., Butterfield D. A., Hayes J. M., Schrenk M. O., Olson E. 42, 377–381. J., Proskurowski G., Jakuba M., Bradley A., Larson B., Riedinger N., Brunner B., Formolo M. J., Solomon E., Kasten S., Ludwig K., Glickson D., Buckman K., Bradley A. S., Brazelton Strasser M. and Ferdelman T. G. (2010) Oxidative sulfur W. J., Roe K., Elend M. J., Delacour A., Bernasconi S. M., cycling in the deep biosphere of the Nankai Trough, Japan. Lilley M. D., Baross J. A., Summons R. E. and Sylva S. P. Geology 38, 851–854. (2005) A serpentinite-hosted ecosystem: the Lost City hydro- Rouxel O., Ono S., Alt J. C., Rumble D. and Ludden J. (2008) thermal field. Science 307, 1428–1434. Sulfur isotope evidence for microbial sulfate reduction in Klein F. and Bach W. (2009) Fe–Ni–Co–O–S phase relations in altered oceanic basalts at ODP Site 801. Earth Planet. Sci. peridotite-seawater interactions. J. Petrol. 50, 37–59. Lett. 268, 110–123. Krouse H. R., Brown H. M. and Farquharson R. B. (1977) Sulfur Sakai H., Casadevall T. J. and Moore J. G. (1982) Chemistry and isotope compositions of sulfides and sulfates, DSDP Leg 37. isotope ratios of sulfur in basalts and volcanic gases at Can. J. Earth Sci. 14, 787–793. Kilauea Volcano, Hawaii. Geochim. Cosmochim. Acta 46, Lasaga A. C., Otake T., Watanabe Y. and Ohmoto H. (2008) 729–738. Anomalous fractionation of sulfur isotopes during heteroge- Sakai H., Des Marais D. J., Ueda a and Moore J. G. (1984) neous reactions. Earth Planet. Sci. Lett. 268, 225–238. Concentrations and isotope ratios of carbon, nitrogen and Lećuyer C. and Ricard Y. (1999) Long-term fluxes and budget of sulfur in ocean-floor basalts. Geochim. Cosmochim. Acta 48, ferric iron: implication for the redox states of the Earth’s mantle 2433–2441. and atmosphere. Earth Planet. Sci. Lett. 165, 197–211. Sasaki A., Arikawa Y. and Folinsbee R. (1979) Kiba reagent Lefticariu L., Pratt L. and Ripley E. (2006) Mineralogic and sulfur method of sulfur extraction applied to isotopic work. Bull. Geol. isotopic effects accompanying oxidation of pyrite in millimolar Surv. Jpn. 30, 241–245. solutions of hydrogen peroxide at temperatures from 4 to 150 Sawyer, D.S., Whitmarsh, R.B., Klaus, A., et al., 1994. Proc. ODP, °C. Geochim. Cosmochim. Acta 70, 4889–4905. Init. Repts, 149: College Station, TX (Ocean Drilling Program). Luz B., Barkan E. and Bender M. L. (1999) Triple-isotope doi:10.2973/odp.proc.ir.149.1994. composition of atmospheric oxygen as a tracer of biosphere Schrenk M. O., Kelley D., Bolton S. a. and Baross J. a. (2004) Low productivity. Nature, 547–550. archaeal diversity linked to subseafloor geochemical processes McCarthy M. D., Beaupre´ S. R., Walker B. D., Voparil I., at the Lost City Hydrothermal Field, Mid-Atlantic Ridge. Guilderson T. P. and Druffel E. R. M. (2010) Chemosynthetic Environ. Microbiol. 6, 1086–1095. origin of 14C-depleted dissolved organic matter in a ridge-flank Schoonen, M.A.A., 2004. Mechanisms of sedimentary pyrite hydrothermal system. Nat. Geosci. 4, 32–66. formation, in: Sulfur biogeochemistry-Past and Present. Geo- McCollom T. and Bach W. (2009) Thermodynamic constraints on logical Society of America Special Paper 379. Geological hydrogen generation during serpentinization of ultramafic Society of America; 1999, pp. 117–134. rocks. Geochim. Cosmochim. Acta 73, 856–875. Shanks W. C. (2001) Stable isotopes in seafloor hydrothermal Mottl M. J., Wheat C. G., Baker E., Becker N., Davis E., Feely R., systems: vent fluids, hydrothermal deposits, hydrothermal Grehan A., Kadko D., Lilley M., Massoth G., Moyer C. and alteration, and microbial processes. Rev. Mineral. Geochem. Sansone F. (1998) Warm springs discovered on 3.5 Ma oceanic 43, 469–525. crust, eastern flank of the Juan de Fuca Ridge. Geology 26, 51. Shanks W. C., Bischoff J. and Rosenbauer R. (1981) Seawater Northrop D. B. (1975) Steady-state analysis of kinetic isotope sulfate reduction and sulfur isotope fractionation in basaltic effects in enzymic reactions. Biochemistry 14, 2644–2651. systems: Interaction of seawater with fayalite and magnetite at Ohmoto H. and Goldhaber M. (1997) Sulfur and carbon isotopes. 200–350 °C. Geochim. Cosmochim. Acta 45, 1977–1995. In Geochemistry of Hydrothermal Ore Deposits, vol. III (ed. H. Shen Y., Farquhar J., Zhang H., Masterson A., Zhang T. and L. Barnes). John Wiley & Sons, New York, pp. 517–611. Wing B. A. (2011) Multiple S-isotopic evidence for episodic 340 S. Ono et al. / Geochimica et Cosmochimica Acta 87 (2012) 323–340

shoaling of anoxic water during Late Permian mass extinction. Wheat C., Jannasch H., Kastner M., Plant J., DeCarlo E. and Nat. Commun. 2, 210. Lebon G. (2004) Venting formation fluids from deep-sea Sim M. S., Bosak T. and Ono S. (2011a) Large sulfur isotope boreholes in a ridge flank setting: ODP Sites 1025 and 1026. fractionation does not require disproportionation. Science 333, Geochem. Geophys. Geosyst. 5, Q08007. http://dx.doi.org/ 74–77. 10.1029/2004GC000710. Sim M. S., Ono S., Donovan K., Templer S. P. and Bosak T. Wheat C. G., Elderfield H., Mottl M. J. and Monnin C. (2000) (2011b) Effect of electron donors on the fractionation of sulfur Chemical composition of basement fluids within an oceanic isotopes by a marine Desulfovibrio sp. Geochim. Cosmochim. ridge flank: implications for along-strike and across-strike Acta 75, 4244–4259. hydrothermal circulation. J. Geophys. Res. 105, 13437-13. Stevens T. O. and McKinley J. P. (1995) Lithoautotrophic Wheat C. G., Mottl M. J. and Rudnicki M. (2002) Trace element microbial ecosystems in deep basalt aquifers. Science 270, 450. and REE composition of a low-temperature ridge-flank hydro- Stevens T. O. and McKinley J. P. (2000) Abiotic controls on H2 thermal spring. Geochim. Cosmochim. Acta 66, 3693–3705. production from basalt-water reactions and implications for Young E. D., Galy A. and Nagahara H. (2002) Kinetic and aquifer biogeochemistry. Environ. Sci. Technol. 34, 826–831. equilibrium mass-dependent isotope fractionation laws in Takai K., Nakamura K., Toki T., Tsunogai U., Miyazaki M., nature and their geochemical and cosmochemical significance. Miyazaki J., Hirayama H., Nakagawa S., Nunoura T. and Geochim. Cosmochim. Acta 66, 1095–1104. Horikoshi K. (2008) Cell proliferation at 122 degrees C and isotopically heavy CH4 production by a hyperthermophilic Associate editor: Edward M. Ripley methanogen under high-pressure cultivation. Proc. Natl. Acad. Sci. U. S. A. 105, 10949–10954.