The role of serpentinites in cycling of carbon and sulfur: Seafloor serpentinization and subduction...

Lithos 178 (2013) 40–54

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Lithos

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Invited review article

The role of serpentinites in cycling of carbon and sulfur: Seafloor serpentinization andsubduction metamorphism

Jeffrey C. Alt a,⁎, Esther M. Schwarzenbach b,1, Gretchen L. Früh-Green b, Wayne C. Shanks III c,Stefano M. Bernasconi b, Carlos J. Garrido d, Laura Crispini e, Laura Gaggero e,José A. Padrón-Navarta f, Claudio Marchesi d

a Dept. Earth and Environmental Sciences, The University of Michigan, Ann Arbor, MI 48109 USAb Dept. Earth Sciences, ETH Zurich, CH-8092 Zurich, Switzerlandc U.S. Geological Survey, 973 Denver Federal Center, Denver, CO 80225 USAd Instituto Andaluz de Ciencias de la Tierra (IACT), CSIC & UGR, Avenida las Palmeras 4, Armilla, 18100 Granada, Spaine DISTAV, University of Genova, Corso Europa 26, 16132 Genova, Italyf Géosciences Montpellier, Univ. Montpellier 2 & CNRS, 34095 Montpellier, France

⁎ Corresponding author. Tel.: +1 7347648380; fax: +E-mail address: [email protected] (J.C. Alt).

1 Now at Virginia Tech Geosciences, 4044 Derring Ha

0024-4937/$ – see front matter © 2012 Elsevier B.V. Allhttp://dx.doi.org/10.1016/j.lithos.2012.12.006

a b s t r a c t
a r t i c l e i n f o
Article history:Received 31 August 2012Accepted 6 December 2012Available online 23 December 2012

Keywords:SerpentiniteCarbonSulfurGeochemical cyclingSubductionStable isotopes

We summarize the uptake of carbon and sulfur during serpentinization of seafloor peridotites, and discuss thefate of these volatiles during subduction of serpentinite. We use a simplified classification to divide seafloorserpentinization into high-temperature and low-temperature processes. High-temperature serpentinizationtypically involves heat andmass transfer from gabbro intrusions, leading to addition of hydrothermal sulfide sul-fur (up to >1 wt.%) having high δ34S values (+5 to +10‰). Total carbon contents of bulk rocks are elevated(0.008–0.603 wt.%) compared to mantle values and δ13CTotal C values of −3‰ to −17.5‰ result from mixturesof organic carbon and seawater-derived carbonate. Low-temperature serpentinization is generally characterizedby microbial reduction of seawater sulfate, which leads to addition of sulfide sulfur (up to 1.4 wt.%) havingnegative δ34S values (down to −45‰), although local closed-system conditions can lead to reservoir effectsand positive δ34S values (up to+27‰). Extensive circulation of cold seawater can cause oxidation, loss of sulfide,and addition of seawater sulfate resulting in high δ34STotal-S values. High total carbon contents (0.006–7.2 wt.%)and δ13C values of −26 to +2.2‰ result from addition of variable proportions of organic carbon andseawater-derived carbonate to serpentinite. We estimate that serpentinization at mid ocean ridges is a sink for0.35–0.64×1011 mol C y−1 and 0.13–1.46×1011 mol S y−1, comparable to the sinks of these elements perunit volume of mafic oceanic crust. Serpentinization in the subducting plate at subduction zones may further af-fect chemical budgets for serpentinization.During subduction metamorphism, sulfur and carbon contents remain unaffected by recrystallization of seafloorlizardite and chrysotile to antigorite, and formation of minor olivine. Dehydration of antigorite-serpentinitesto chlorite–harzburgites at higher pressure and temperature results in loss of 5 wt.% water, and an average of260 ppmsulfur is lost as sulfate having δ34S=14.5‰, whereas carbon is unaffected. These volatiles can inducemelt-ing and contribute to 34S enrichments and oxidation of the sub-arcmantlewedge. Serpentinized oceanic peridotitescarry isotopically fractionatedwater, carbon and sulfur into subduction zones. Up to 0.49×1011 mol sulfur y−1 and1.7×1011 mol carbon y−1 are subducted in serpentinites, less than 3% of the total subduction budgets for each ofthese elements. Isotopically fractionated carbon, sulfur, and water remain in serpentinite dehydration products,however, and can be recycled deeper into the mantle where they may be significant for volatile budgets of thedeep Earth.

© 2012 Elsevier B.V. All rights reserved.

Contents

1. Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 412. Oceanic serpentinization . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 42

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41J.C. Alt et al. / Lithos 178 (2013) 40–54

2.1. Low-temperature serpentinization . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 422.1.1. Iberian margin . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 422.1.2. Northern Apennine ophiolites . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 432.1.3. The Lost City Hydrothermal Field . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 432.1.4. 15° 20′ N Fracture Zone, MAR . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 44

2.2. High-temperature serpentinization . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 442.2.1. The MARK area . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 452.2.2. 15° 20′N Fracture Zone, MAR . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 452.2.3. Northern Apennine ophiolites . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 452.2.4. Atlantis Massif, MAR . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 462.2.5. Hess Deep . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 46

3. Abundance of serpentinite . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 463.1. Abundance of serpentinite in oceanic basement . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 463.2. Abundance of serpentinite in subducting slabs . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 47

4. Uptake of carbon and sulfur in seafloor serpentinites . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 484.1. Carbon . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 484.2. Sulfur . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 49

5. Effects of subduction metamorphism on carbon and sulfur in serpentinites . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 495.1. Recrystallization of lizardite and crysotile to antigorite . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 495.2. Dehydration of antigorite serpentinites . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 505.3. Recycling of carbon and sulfur during subduction of serpentinite . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 50

5.3.1. Sulfur . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 505.3.2. Carbon . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 51

6. Conclusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 51Acknowledgments . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 51References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 51

1. Introduction

Serpentinized peridotite is now recognized as an important com-ponent of oceanic basement. Mantle peridotite occurs at the seaflooron slow spreading ridges in areas of tectonic extension, where ocean-ic core complexes are exposed by unroofing of gabbroic intrusivesand upper mantle along detachment faults (e.g., Blackman et al.,1998; Cann et al., 1997; Cannat et al., 1995a, 2010; Dick et al., 2008;Smith et al., 2006; Tucholke and Lin, 1994). Peridotite is also commonon ultraslow spreading ridges where mantle material may be exposedat the seafloor symmetrically along the spreading axes (Cannat etal., 2006; Dick et al., 2003; Michael et al., 2003). In crust formed at

Limit of fluid circul

High-T Serpe

Detachment Fault

Hydrotherma+ sulfide depo

Melt orweak zone

Mafic upper crust

Fig. 1. Schematic diagram illustrating high- and low-temperature serpentinzation processeswhich roots in melt or weak zone at depth, marking the limit of fluid penetration. High-tempinto footwall as mantle material is uplifted (white arrows indicate deepening high-temperagabbro intrusions, and results in further high-temperature serpentinization (green shading(light shading) is driven by cooling of the lithosphere and is associated with faulting, fracttemperature serpentinization, in unaltered peridotite, or as cool seawater is entrained into h2012b; Andreani et al., 2007; McCaig et al., 2007, 2010; Petersen et al., 2009; Tucholke et al., 2

intermediate and fast spreading rates peridotite is exposed locally atpropagating rifts and in fracture zones (Früh-Green et al., 1996;Hekinian et al., 1996). Peridotites are readily serpentinized overa wide range of conditions, and serpentinization is important as asink for water, carbon, sulfur, chlorine, boron, arsenic, and nitrogen(e.g., Alt and Shanks, 1998, 2003; Alt et al., 2007; Bach et al., 2004;Barnes and Sharp, 2006; Bonatti et al., 1984; Boschi et al., 2006, 2008;Delacour et al., 2008a,b,c; Deschamps et al., 2010; Früh-Green et al.,2004; Halama et al., 2012; Kendrick et al., 2011; Paulick et al., 2006;Scambelluri et al., 2004; Schwarzenbach, 2011; Schwarzenbach et al.,2012; Vils et al., 2011). Detachment faults and associated faults andshear zones can focus fluid flow, leading to serpentinization at high

ation

Gabbro Intrusions

Low-Tserpentinization

ntinization

l ventssits

in relation to detachment faulting. Fluid circulation is focused along detachment fault,erature serpentinization proceeds as fluids penetrate into footwall, and reaches deeperture serpentinization front). Hydrothermal fluid circulation (black arrows) is driven by) and formation of sulfide deposits at seafloor vents. Low temperature serpentinizationuring and exposure of peridotite at the seafloor. This can occur superimposed on highydrothermal upflow zones beneath seafloor vents and sulfide deposits (After Alt et al.,008).

90˚W 0˚

20˚

0˚

60˚

N

40˚Rainbow

LogatchevMARK

Lost CityAtlantis massif

15˚20’ FZ

Hess Deep

Almirez

IberianMargin

VoltriApennine

Fig. 2. Map showing locations of sites discussed in the text.

42 J.C. Alt et al. / Lithos 178 (2013) 40–54

temperatures at various depths, down to the base of the crust (Fig. 1;Andreani et al., 2007; Boschi et al., 2006; McCaig et al., 2007, 2010;Schroeder and John, 2004). Gabbro intrusions provide heat that candrive high-temperature ultramafic-hosted hydrothermal systems thatresult in black-smoker type vents and sulfide deposits atop serpentiniteon the seafloor, such as the Logatchev and Rainbow sites on themid-Atlantic Ridge (Charlou et al., 2002; Douville et al., 2002; McCaigand Harris, 2012; McCaig et al., 2010; Petersen et al., 2009;Schmidt et al., 2007; Seyfried et al., 2004, 2010). Lower-temperatureserpentinization can occur as seawaterfluids circulate through fracturesand faults near the seafloor, driven by the general cooling of the litho-sphere and exothermic serpentinization reactions (Agrinier et al.,1988; Allen and Seyfried, 2004; Alt et al., 2007; Früh-Green et al.,2003; Kelley et al., 2001; McCaig, et al., 2007; Schwarzenbach, 2011;Schwarzenbach et al., 2012).

Serpentinization associated with subduction zones is also importantfor cycling of volatiles and fluid mobile elements. Serpentinizationof suboceanic mantle can occur outboard of subduction zones, whereflexure of the lithosphere allows faulting and penetration of fluids tohydrate themantle (Peacock, 2001; Ranero and Sallares, 2004). In addi-tion, the mantle wedge in subduction zones is serpentinized by fluidsreleased from the slab, and can be dragged downward to releasefluids and various elements at greater depths (Deschamps et al., 2010;Hattori and Guillot, 2003; Savov et al., 2005; Tatsumi et al., 1986).Serpentinization and other hydration reactions in the subduction chan-nel itself also form phases that can carry water and volatiles deep intosubduction zones (Bebout and Barton, 2002; Padrón-Navarta et al.,2010; Spandler et al., 2008).

Serpentine is stable to high pressure and temperature, and cancarry significant amounts of water to depths of up to 200 km in sub-duction zones, where dehydration releases volatiles that can triggermelting in the mantle wedge (Hacker, 2008; Rüpke et al., 2004; Ulmerand Trommsdorff, 1995; Wallmann, 2001). Seafloor serpentinitesare enriched in carbon and sulfur compared to mantle values, and,depending on subduction pathways, serpentinite and its dehydrationproducts can release these volatiles to the mantle wedge or recyclethem into the mantle (Alt et al., 2012a; Kerrick and Connolly, 1998). Inthis paper, we summarize data for carbon and sulfur in serpentinitesfrom the seafloor and from ophiolites, and evaluate the role of oceanicserpentinites in the uptake of carbon and sulfur during serpentinizationon the seafloor. We also assess the effects of subduction metamorphismon carbon and sulfur in oceanic serpentinites that were subductedto varying depths and temperatures.

2. Oceanic serpentinization

Serpentinization of oceanic peridotite occurs over a wide range offluid–rock ratios, temperatures, redox, and fluid compositions (Altand Shanks, 1998, 2003; Andreani et al., 2007; Bach et al., 2004,2006; Boschi et al., 2008; Delacour et al., 2008a,b,c; Deschamps et al.,2010; Frost and Beard, 2007; Früh-Green et al., 1996, 2004; Klein etal., 2009; Schwarzenbach, 2011; Schwarzenbach et al., 2012;Seyfried et al., 2004). These conditions commonly evolve as ser-pentinization reactions proceed, leading to superimposed effects. Insome cases one geochemical tracer may record high-temperatures,whereas another tracer may show the effects of low temperatureprocesses. For example, amphibole and talc replacing orthopyroxeneor 18O-depletion of serpentinites can indicate the effects of hightemperature processes (>350 °C), whereas sulfur isotope data for thesame rocks can indicate the effects of later microbial reactions at lowtemperatures (b120 °C; Alt and Shanks, 1998; Früh-Green et al.,1996; Schwarzenbach et al., 2012). Carbonate veins can form fromhighly reacted hydrothermal fluids early during serpentinization at rel-atively high temperatures (150–350 °C), but they can also form laterin the same rocks from seawater at low temperatures (0–40 °C; Altand Shanks, 2003; Bach et al., 2011; Früh-Green et al., 2003; Klein and

Garrido, 2011; Schwarzenbach, 2011). With these caveats, we usea simplified but convenient classification to discuss carbon andsulfur in serpentinites, and divide seafloor serpentinization into“high-temperature” and “low-temperature” processes (Alt et al., 2007,2012b). Locations of sites discussed in this paper are shown in Fig. 2.

2.1. Low-temperature serpentinization

Low-temperature serpentinization occurs where fluid flow isdriven by the general cooling of the lithosphere and exothermicserpentinization reactions, and is influenced by faulting, fracturing,and exposure of peridotite to seawater at the seafloor. In the followingSections 2.1.1–2.1.4 we show that serpentinites affected by low-temperature processes contain variablemixtures of seawater carbonateand organic carbon, and can contain carbonate veins or may becarbonate-cemented breccias (ophicalcites) having high carbon con-tents (several wt.%) and seawater δ13C values. The rocks have high sul-fur contents compared to mantle peridotites, and have generallynegative δ34S values as the result of microbial activity. The presence ofmicrobial effects indicates temperatures less than 120 °C, which arecommonly borne out by oxygen isotope measurements on bulk rocksand mineral separates. Locally, extensive interaction of serpentiniteswith cold seawater at the seafloor or along shallow faults can lead tooxidation of sulfide minerals, loss of sulfide, and incorporation of sea-water sulfate.

Variable amounts of sulfate are common in seafloor serpentinites.The sulfate is in part derived from seawater, which has a high δ34Svalue (modern seawater=+21‰), but is also derived from oxidationof sulfide minerals, which have lower δ34S values. Moreover, bulkrock sulfate-sulfur can have lower δ34S than associated sulfide min-erals in the rocks, which is in the opposite sense for equilibrium,where the oxidized species is enriched in 34S relative to the reducedspecies. This has been attributed to a kinetic isotope effect during ox-idation of sulfide minerals in the rocks (Alford et al., 2011; Alt andShanks, 1998, 2003; Alt et al., 2007, 2012a, 2012b; Delacour et al.,2008a, 2008b; Schwarzenbach et al., 2012).

2.1.1. Iberian marginPeridotites at the Iberian margin were exposed at the seafloor

by tectonic extension during opening of the Atlantic Ocean basinin the early Cretaceous, and several drill holes of the Ocean DrillingProgram (ODP) penetrate up to 153 m into the serpentinite base-ment (Whitmarsh and Sawyer, 1996). Bulk rock 18O enrichments


(δ18O≈10‰) and large serpentine-magnetite oxygen isotope fraction-ations (12‰) indicate that serpentinization occurred predominantly atlow temperatures, b200 to b50 °C (Agrinier et al., 1988, 1996). Therocks have high sulfide-sulfur contents and generally negative δ34Svalues (up to 0.4 wt.% and to−45‰, respectively, compared to mantlevalues of 100–250 ppm and ~0.1‰; Fig. 3, Table 1; Alt and Shanks,1998, 2003; Sakai et al., 1984; Schwarzenbach et al., 2012). In thedeepest cores (100–150 m subbasement) δ34S values exhibit a progres-sive increasewith depth as the result of reservoir effects in a closed sys-tem, leading to high δ34S values in the deepest samples (up to +27‰at 140–150 m; Fig. 3). At the low temperatures of serpentinization, mi-crobial activity is the only reasonable mechanism to account for sulfidegain and the negative δ34S values, and Alt and Shanks (1998) suggestedthat microbial communities were supported by hydrogen generatedduring serpentinization reactions.

Serpentinite contains abundant carbonate veins in the upper 30–70 mof these sections, and is partly replaced by red smectite and carbonate,which formed from seawater at low temperatures, 10–30 °C (Agrinieret al., 1988, 1996; Milliken and Morgan, 1996; Schwarzenbach, 2011).Compared to mantle peridotites (30 ppm C, δ13C ~−5‰; Dasguptaand Hirschman, 2010), the rocks are enriched in total carbon (up to9.6 wt.%) although the deeper serpentinites, down to 153 m, lack visi-ble carbonate and have lower total carbon contents (b0.5 wt.%;Schwarzenbach, 2011).

Carbon in the bulk rocks comprisesmixtures of seawater carbonateand organic carbon, and the rocks span the entire field of data plottedin Fig. 4. The high-carbon samples are dominated by seawater carbon-ate, whereas the low-carbon samples contain greater proportions oforganic carbon and have negative δ13C values (down to −28‰;Schwarzenbach, 2011).

2.1.2. Northern Apennine ophiolitesThe northern Apennine ophiolites of the internal Ligurides are

fragments of Jurassic oceanic crust that formed at slow or ultra-slowspreading centers, on the oceanward side of an ocean–continent

Beigua

-50

-40

-30

-20

-10

10

20

30

40

0

10 100 1000 10000Sulfur content (ppm)

open systemmicrobial sufate reduction

closed system microbialsulfate reduction

hydrothermal sulfide

oxidation

melting

Serpentinite Chl-harz

Almirez:

δ34S

(‰

VC

DT

)

Fig. 3. Bulk rock δ34S vs sulfur content of serpentinites. Mantle is indicated by rectangle,arrows and grey fields indicate processes and effects during serpentinization of seafloorperidotites (after Alt et al., 2007). Hydrothermal sulfide is added to peridotite duringserpentinization by high-temperature (350–400 °C) hydrothermal fluids (dotted linesand medium grey field). Microbial reduction of seawater sulfate at low temperatures re-sults in negative δ 34S values in an open system (grey arrows and light grey field), andin a closed system can lead to reservoir effects and 34S-enrichment (dashed grey arrowand light greyfields). Oxidation results in loss of sulfurwith no fractionation orwith kinet-ic isotope fractionation and 34S-enrichment of residual sulfide (dark grey field). Meltingcauses loss of sulfur with no isotope fractionation. Data fields include sulfide-sulfur datafor seafloor serpentinites (Alt and Shanks, 1998, 2003; Alt et al., 2007; Delacour et al.,2008a,b) and total sulfur data for Apennine ophiolite serpentinites (Alt et al., 2012b;Schwarzenbach et al., 2012). Symbols indicate total sulfur in high-pressure antigoriteserpentinites and chlorite–harzburgite dehydration products (from Alt et al., 2012a,b).

transition (Piccardo, 2008; Principi et al., 2004). There is someobduction-related thrust faulting and folding, but the structure ofthe ophiolites is essentially identical to that of oceanic core com-plexes exposed by detachment faulting at mid ocean ridges. Upliftedmantle peridotites were intruded by MORB gabbro bodies and dikes,exposed at the seafloor and serpentinized, and then covered withpillow lavas and pelagic sediments (Lagabrielle and Cannat, 1990;Menna, 2009; Piccardo, 2008; Principi et al., 2004; Scambelluri et al.,1997).

Schwarzenbach (2011) and Schwarzenbach et al. (2012) presentresults for serpentinites from quarries in the northern Apennines thatrange from intact (unveined) serpentinite, to rocks veined by carbon-ate, to those that are brecciated, cemented and replaced by carbonate(ophicalcites). These are all from within ~30 m of the paleo-seafloor,and were serpentinized mainly at moderate to low temperatures(~240° to b150 °C) with carbonates precipitating at temperatures of150 °C down to 50 °C (Schwarzenbach, 2011; Schwarzenbach et al.,2012). The rocks are enriched in sulfur and have generally negativeδ34S values (up to 1.4 wt.% and to –34‰, respectively; Fig. 3; Table 1),as the result of microbial reduction of seawater sulfate. A few rocks ex-hibit 34S-enrichments that result from reservoir effects duringclosed-system microbial sulfate reduction (Fig. 3). Samples of localamphibole-rich shear zones record early high-temperature processes,and are enriched in 34S and copper from hydrothermal fluids(Schwarzenbach et al., 2012).

Like the seafloor serpentinites, the bulk ophiolitic serpentinitescontain mixtures of carbon comprising seawater carbonate and a com-ponent of organic carbon (Fig. 4). Total carbon contents range up to7.25 wt.% and δ13C of total carbon ranges from −16.4‰ to +2.2‰(Table 1; Schwarzenbach, 2011).

2.1.3. The Lost City Hydrothermal FieldThe Atlantis Massif near the Mid-Atlantic Ridge (MAR) is an

oceanic core complex exposed by detachment faulting (Fig. 2;Blackman et al., 1998; Ildefonse et al., 2007). Serpentinized perido-tites from the off-axis Lost City hydrothermal field (LCHF), on thesouthern wall of the Atlantis Massif, exhibit the effects of evolvingprocesses over a range of temperatures. Although a large gabbrobody is present at the core of the massif (Ildefonse et al., 2007),there is no indication of a gabbroic heat source or chemical effectsof gabbro in hydrothermal fluids venting at Lost City (Allen andSeyfried, 2004; Foustoukos et al., 2008; Kelley et al., 2001). Fluidsvent from serpentinite at 40–90 °C, and the fluid compositions indi-cate relatively low-temperature (b200 °C) serpentinization reactions(Allen and Seyfried, 2004; Foustoukos et al., 2008; Kelley et al., 2005;Proskurowski et al., 2008). Fluid circulation is driven by the generalcooling of the lithosphere and by exothermic serpentinization reac-tions (Allen and Seyfried, 2004; Foustoukos et al., 2008; Kelley et al.,2001; Lowell, 2010; Lowell and Rona, 2002). Sulfur in serpentinitesassociated with active venting reflects the influence of microbial ac-tivity, with high sulfide-sulfur contents (590–4560 ppm) and highlyvariable δ34S values (−22.9 to +27.0‰) that reflect microbial sulfatereduction under conditions that ranged from open- to closed-system(Fig. 3, Table 1; Delacour et al., 2008a).

The shear zone that makes up the detachment fault surface of theAtlantis Massif was hydrothermally altered over a range of tempera-tures, with initial hydration of mafic rocks at temperatures above600 °C, formation of talc and amphibole schists in the detachmentshear zone at 270–350 °C, and serpentinization of the footwall peri-dotite predominantly at temperatures of ~150–250 °C (Boschi et al.,2008; Schroeder and John, 2004). Serpentinized and altered perido-tites from the southern wall of the massif have low sulfide-sulfur con-tents and high δ34Ssulfide values (7.7–14.9‰; Fig. 3, Table 1). These areinterpreted to reflect serpentinization at high water/rock ratios, witha kinetic isotope fractionation during oxidation of primary sulfide

Table 1Summary of concentration and isotope data for C and S in serpentinite, for areas discussed in text.

Low temperature serpentinites High temperature serpentinites High-pressure rocks

1520Hole1272A &1274

NorthernApennine

IberianMargin

Lost CityHydroth.field

1520FZHole1268A

1520FZSite1270

1520FZSite1271

MARK NorthernApennine

Hessdeep

Atlantismassif

Voltri(Beigua)

AlmirezSerpentinite

AlmirezChloriteharzburgite

Totalcarbonwt.%

Mean 0.319 1.063 1.718 0.180 0.027 0.125 0.109 0.120 0.033 0.079 0.067 0.028 0.051 0.050StDev 0.693 1.631 2.453 0.365 0.012 0.195 0.127 0.057 0.007 0.076 0.072 0.004 0.036 0.023Median 0.095 0.352 0.302 0.040 0.025 0.057 0.071 0.112 0.034 0.049 0.046 0.028 0.041 0.044Min 0.070 0.014 0.033 0.006 0.008 0.025 0.030 0.002 0.026 0.019 0.015 0.023 0.018 0.019Max 2.540 7.247 9.629 1.601 0.055 0.603 0.420 0.260 0.048 0.251 0.271 0.032 0.128 0.083

δ13C totalcarbon

Mean −4.6 −9.3 −13.7 −8.3 −6.0 −11.9 −9.3 −15.4 −12.9StDev 5.1 8.6 9.6 1.5 3.6 1.1 4.5 2.0Median −2.0 −6.8 −13.6 −8.0 −11.6 −9.0 −16.1 −12.8Min −16.4 −25.3 −26.0 −10.4 −17.5 −10.9 −20.2 −15.8Max 2.2 1.6 0.5 −6.0 −3.2 −8.5 −9.7 −9.6

Sulfidesulfurppm

Mean 131 929 818 535 2580 1270 587 36 1390 645 785 415StDev 112 2450 986 1295 3140 2120 508 40 2384 619 664 426Median 120 320 453 16 1179 665 404 20 313 581 892 279Min 31 7 1 10 460 0 0 26 136 1 0 103 29 46Max 386 14,562 4320 4665 9421 19 420 9524 1266 165 8248 1316 2002 1484

Total sulfurppm

Mean 875 939 2328 858 3846 825 1755 711 321 1421 689 914 652StDev 468 2390 2733 1436 4504 530 2082 546 216 2223 623 733 623Median 787 370 1230 273 1721 640 1325 574 332 642 635 1079 596Min 415 20 140 25 1070 217 420 414 136 7 236 133 100 145Max 1976 14,562 13,320 5397 13,626 728 1600 9986 1437 857 8443 1353 2309 2311

δ34Ssulfide-S

Mean −16.8 −8.6 −20.2 9.0 8.0 7.4 5.7 −5.8 4.6 10.7 5.5 4.4StDev 14.4 8.9 18.0 12.0 2.3 2.7 3.0 8.7 5.4 2.5 5.3 4.7Median −16.7 −7.9 −19.6 10.9 8.1 7.4 4.4 −2.9 4.9 10.8 7.1 5.0Min −32.1 −34.4 −45.1 −22.9 4.4 3.8 3.1 −23.7 −1.3 6.9 −4.6 −5.1Max −1.5 18.9 27.2 27.1 10.8 4.0 12.7 11.2 4.9 13.9 14.3 11.2 10.2

δ34S total S Mean 1.5 −8.5 −8.5 15.1 6.9 11.1 4.5 9.0 4.4 9.3 5.0 1.2StDev 2.2 8.0 10.4 8.1 2.5 4.3 3.2 6.2 5.0 2.3 4.1 3.9Median 2.1 −8.4 −7.6 18.3 7.5 11.1 3.7 8.9 3.9 9.8 5.7 2.7Min −1.5 −24.6 −29.7 −8.4 3.5 4.0 4.2 1.3 −3.3 −0.8 6.0 −3.5 −6.6Max 3.3 15.8 14.0 20.2 9.9 11.3 7.8 19.4 9.8 18.0 12.1 11.5 9.7 4.6

SO4/Total S Mean 0.877 0.189 0.465 0.750 0.351 0.960 0.908 0.466 0.204 0.867 0.205 0.142 0.235 0.414Stdev 0.112 0.234 0.337 0.307 0.101 0.284 0.162 0.121 0.230 0.118 0.213 0.188

Data source 1,2 3,4 3,4,5 6,7,15 1,2 1,2 1,2 8,9,13,15 10 5,11,12,13 6,7 10 14 14

1. Kelemen et al. (2004); 2. Alt et al. (2007); 3. Schwarzenbach (2011); 4. Schwarzenbach et al. (2012); 5. Alt and Shanks (1998); 6. Delacour et al. (2008a);7. Delacour et al.(2008b); 8. Alt and Shanks (2003); 9. Cannat et al. (1995b); 10. Alt et al. (2012b); 11. Früh-Green et al. (1996); 12. Gillis et al. (1993); 13. Früh-Green et al. (2004); 14. Alt etal. (2012a); 15. Delacour et al. (2008c).


and loss of sulfate resulting in 34S-enrichment of the residual sulfur(Delacour et al., 2008a).

Like the other low-temperature serpentinites, the bulk carboncontents and δ13C values of LCHF serpentinites fall along a mixingline between seawater carbonate and organic carbon (Fig. 4, Table 1;Delacour et al., 2008c). Samples containing carbonate veins havehigh carbon contents (up to 1.6 wt.%) and seawater-like δ13C values,whereas samples lacking veins have low C contents b0.1 wt.%) andnegative δ13C values (to−18‰). Some carbonate veins in the basementhave δ13C values as low as−6‰, suggesting oxidation of a 13C-depletedcomponent, such asmethane, derived from fluids at depth (Früh-Greenet al., 2003). Some of these carbonates have low δ18O values that indi-cate precipitation temperatures of ~50 to 225 °C. Carbonate in chim-neys, breccias, and sediment at LCHF has positive δ13C values (1.4 to3.2‰), consistent with fractionation of some 12C into methane or hy-drocarbons in hydrothermal fluids (Früh-Green et al., 2003).

2.1.4. 15° 20′ N Fracture Zone, MARPeridotite and gabbro are exposed by detachment faulting around

the 15° 20′ fracture zone on the MAR, and ODP Holes 1272A and1274A near the intersections of the fracture zone and the MAR pene-trate 131 and 156 meters, respectively, into serpentinized peridotite(Kelemen et al., 2004). Generally high bulk rock δ18O values (up to8.1%) indicate low-temperature (b150 °C) serpentinization (Alt et al.,2007), but oxygen isotope analyses of serpentine-magnetite mineral

pairs suggest temperatures of 160 °C to >1000 °C (Barnes et al.,2009a). Temperatures above 500 °C estimated from serpentine–magnetite pairs are inconsistent with petrographic observations,however, indicating disequilibrium between serpentine and magne-tite in at least some samples (Barnes et al., 2009a). The high sulfide-Scontents and negative δ34Ssulfide values (up to 3000 ppm and to−32.1%, respectively) are interpreted to reflect microbial reductionof seawater sulfate andare consistentwith thebulk rock 18Oenrichmentsand alteration predominantly at low temperatures (Alt et al., 2007;Table 1; Fig. 3). Both of the drill holes penetrate fault zones, and flow ofcold seawater along these zones led to oxidation of sulfide mineralsand incorporation of seawater sulfate (δ34S=+21‰), resulting in highSO4/ΣS values (0.6–1.0) and δ34STotal S values (up to +3.3‰; Table 1).

Most of the serpentinites contain 0.02–0.16 wt.% total carbon,although those from the upper 50 m in Hole 1272A are veined bycalcium carbonate and contain greater amounts of carbon (up to2.5 wt.%; Table 1; Kelemen et al., 2004). Aragonite veins in Hole1274A have δ13C values of 0–3‰ and formed during circulation ofcold (~0–15 °C) seawater near the seafloor (Bach et al., 2011).

2.2. High-temperature serpentinization

High temperature serpentinization processes commonly reflectthe influence of heat and mass transfer from associated gabbroicintrusions that drive hydrothermal circulation in ultramafic rocks

0

-10

-30

-20

Total carbon content (wt%)2 4 6 8

Beigua

Serpentinite Chl-harz

Almirez:

Serpentinites Mixing

δ13C

(‰

VP

DB

)

Fig. 4. Bulk rock δ13C vs Total carbon content for serpentinites. Data plot along a mixingline between seawater carbonate (δ13C=0‰) and a low-δ13C reduced carbon compo-nent. Dashed line is mixing line between 10 wt.% seawater carbon (0‰) and0.120 wt.% (120 ppm) reduced carbon (−25‰). Shaded field encompasses 212 datapoints for high- and low-temperature serpentinites from the seafloor and from Apen-nine ophiolites (Alt et al., 2012a, 2012b; Delacour et al., 2008c; Schwarzenbach, 2011).Plotted data points are for high-pressure antigorite serpentinites and chlorite–harzburgite dehydration products (from Alt et al., 2012a,b).


at high temperatures, ~350–400 °C. These hydrothermal systems re-sult in seafloor vents and sulfide deposits like those at the Rainbowand Logatchev sites on the MAR (Figs. 1 and 2). The resultingserpentinites tend to be depleted in 18O compared to mantle perido-tites, indicating temperatures above 200 °C (Alt et al., 2007; Saccociaet al., 2009). In extreme cases, the serpentinites can also be enrichedin light rare-earth elements and have positive europium anomalies,similar to the vent fluids, and can exhibit the effects of silicametasoma-tism (replacement by talc; Bach et al., 2004; Paulick et al., 2006). Asshown in Sections 2.2.1–2.2.5 below, the rocks are slightly enrichedin carbon compared to fresh mantle peridotite, have negative total car-bon δ13C values, have high sulfur contents and elevated δ34S values(Fig. 3, Table 1). Similar effects can occur along shear zones associatedwith detachment faulting, which can focus flow of hydrothermal fluids(e.g., Boschi et al., 2006; McCaig and Harris, 2012; McCaig et al., 2007;Schwarzenbach et al., 2012). Serpentinization at high temperaturescan also occur where seawater fluids penetrate into hot peridotite, inthe absence of a gabbroic influence (Andreani et al., 2007; Früh-Greenet al., 1996; McCaig et al., 2010).

2.2.1. The MARK areaTwo ocean drilling holes in the MARK area, at the eastern intersec-

tion of the Kane Fracture Zone and the Mid-Atlantic Ridge (MAR),penetrate 126 m and 200.8 m into serpentinized peridotite with in-terspersed thin zones of gabbro in an oceanic core complex (Fig. 2;Cannat et al., 1995a,b). The serpentinites have high sulfur contentsand δ34S values (up to 1 wt.% and 12.7‰, respectively; Table 1) thatcannot be produced by reaction of seawater with peridotite alone(Alt and Shanks, 2003). Incremental reactionmodeling indicates amul-tistage process involving reduction of seawater sulfate and leaching ofsulfide from gabbro during reactions with seawater at higher tempera-tures (>350 °C) where sulfide minerals are soluble, and subsequentdeposition of sulfides during serpentinization of peridotite at lowertemperatures, b350 °C, as the solubility of metal sulfides decreases(Alt and Shanks, 2003). The serpentinites have low δ18O values (2.6–3.7‰; Agrinier and Cannat, 1997), and contain 20–2600 ppm totalcarbon , which is a mixture of seawater carbonate and organic carbonhaving δ13C around −26‰ (Cannat et al., 1995b; Delacour et al.,2008c; Früh-Green et al., 2004). Carbonate veins formed at low temper-atures (near 0 °C) from seawater, and from hydrothermal fluids attemperatures up to >250 °C (Alt and Shanks, 2003). High δ13C ofthe higher-temperature carbonates (up to +4.5‰) result from reduc-tion of seawater carbonate to methane or other hydrocarbons duringserpentinization.

2.2.2. 15° 20′N Fracture Zone, MARODP Sites 1268 and 1270 are located in core complexes ~40 km

south of the 15° 20′ Fracture Zone on the MAR, and both are within~20 km of the high-temperature Logatchev ultramafic-hosted hydro-thermal vent site (Fig. 2; Kelemen et al., 2004). Hole 1268A pene-trates 147.6 m into peridotite that was first serpentinized and thenunderwent a second alteration stage of talc metasomatism throughinteraction with hydrothermal fluids derived from seawater interac-tion with gabbro+peridotite (Alt et al., 2007; Bach et al., 2004;Paulick et al., 2006). Four shallow holes (18–57 m deep) at Site1270 penetrate serpentinized peridotite plus minor hydrothermallyaltered gabbroic rocks. The altered peridotites from the two drill sitesare generally depleted in 18O, having δ18O values down to 2.6‰, al-though a few shallow samples are enriched in 18O consistent withlate stage carbonate veins and low-temperature effects near the sea-floor (Alt et al., 2007; Kelemen et al., 2004).

The rocks from Hole 1268A are enriched in sulfur, with thetalc-altered rocks having the highest sulfur contents (up to 1.2 wt.%,Table 1). The bulk rocks and vein pyrites have high δ34S values(4.4–10.8‰), although a few negative values (to −15.2‰) are

consistent with local late low-temperature effects (microbial sulfatereduction; Alt et al., 2007).

Sulfur in the shallow rocks from Site 1270 and from Site 1271 atthe segment end to the north is highly oxidized by late seawater ef-fects, with high SO4/ΣS values (mostly 0.9–1.0; Table 1). The lowδ18O values for Site 1270 and the high δ34S of sulfide, when present,at Site 1271 indicate high temperature serpentinization processes(Alt et al., 2007), but superimposed cold seawater alteration has af-fected sulfur and carbon in the rocks.

Early calcite veins associated with detachment faulting at Site1271 have high δ13C values (to +8.7‰) reflecting reduction of sea-water carbonate and fractionation of 12C into methane (Bach et al.,2011). These calcites formed from hydrothermal fluids similar tothose venting from the nearby Logatchev hydrothermal field, whereaslater aragonite veins have δ13C values of around 0‰ and formed duringcirculation of cold (~0–15 °C) seawater near the seafloor (Bach et al.,2011). The hydrothermally altered peridotites from Hole 1268A aver-age 0.03 wt.% carbon, but the other sites have higher carbon contents,up to 0.60 wt.%, as the result of later carbonate formation at low tem-peratures (Kelemen et al., 2004; Table 1).

2.2.3. Northern Apennine ophiolitesPeridotites associated with a gabbro intrusion and with seafloor

hydrothermal sulfide deposits in northern Apennine ophiolites wereserpentinized by hydrothermal fluids at elevated temperatures (Altet al., 2012b). Compared to mantle peridotites the serpentinites areenriched in sulfur and have elevated δ34S values (up to 1440 ppmand +9.8‰, respectively; Table 1). These effects occurred in seafloorultramafic-hosted hydrothermal systems driven by mafic intrusions,like those that feed hydrothermal vents at the Logatchev and Rain-bow hydrothermal sites on the seafloor (Alt et al., 2012b). Local neg-ative δ34S values of serpentinites beneath seafloor sulfide depositsin the ophiolite result from low-temperature alteration and microbi-al sulfate reduction during entrainment of cold seawater into theshallow upflow zone beneath hydrothermal vents at the seafloor


(Fig. 1; Alt et al., 2012b). The rocks lack visible carbonate veins andbulk rock data fall at low carbon contents and negative δ13CTotal C valuesalong the mixing line between reduced (organic) carbon and seawatercarbonate (Fig. 4; Table 1; Alt et al., 2012b).

2.2.4. Atlantis Massif, MARIODP Hole 1309D, in the central dome of the Atlantis Massif on

the MAR, penetrates 1.4 km of gabbro with interspersed zones ofserpentinized peridotite (Blackman et al., 2006). The latter wereserpentinized at high temperatures, with the rocks affected by masstransfer through hydrothermal reactions with associated gabbroicrocks (Delacour et al., 2008b). Serpentinized harzburgites and troctolitesare generally enriched in S and 34S (up to 8440 ppm and +12.1‰,respectively; Table 1; Fig. 3). Total carbon contents and δ13C valuesare variable, reflecting mixtures of seawater carbonate and organiccarbon having δ13C=−27‰ (Table 1; Fig. 4; Delacour et al., 2008c).Organic carbon is derived from a combination of sources, includingseawater DOC, abiotic synthesis of hydrocarbons, and microbial activity(Delacour et al., 2008c).

2.2.5. Hess DeepPeridotites at Hess Deep, in the eastern Pacific, were exposed

through rifting as the Galapagos spreading center propagatedwestward. Ocean drilling penetrated 10–94 m into serpentinizedperidotites that record early high-temperature serpentinizationand later superimposed low-temperature effects. Serpentinizationat high temperatures (200–400 °C) and low fluid/rock ratios isindicated by low δ18O values of serpentine mineral separates(3.2–4.1‰) and oxygen isotope values of serpentine–magnetite min-eral pairs (Früh-Green et al., 1996). Serpentinization occurred duringrifting as seawater fluids penetrated into hot peridotite, and there isno evidence for interactionwith gabbro-influenced hydrothermal fluids(Früh-Green et al., 1996). Low-water–rock ratios during initialhigh-temperature serpentinization reactions resulted in little changein sulfur contents (Table 1), and later addition of microbially-derivedsulfide at low temperatures produced negative δ34S values forsulfide-sulfur (to −23.7‰; Table 1; Alt and Shanks, 1998). The drillholes penetrate highly fractured, fault-exposed peridotite, and late oxi-dation by seawater near the seafloor and incorporation of seawatersulfate (+21‰) resulted in high SO4/ΣS values and consequenthigh δ34STotal S values (−3.3 to +18.0‰; Table 1). The rocks contain190–1150 ppm carbon, which has δ13C values of around −6‰ andconsists of mixtures of organic carbon, seawater carbonate, and pos-sibly magmatic carbon (Früh-Green et al., 1996, 2004; Gillis et al.,1993). Carbonate veins in the rocks formed during exposure to sea-water at low temperatures (Blusztajn and Hart, 1996; Früh-Greenet al., 1996).

3. Abundance of serpentinite

3.1. Abundance of serpentinite in oceanic basement

A critical factor in understanding the role of oceanic serpentinitesfor chemical cycling is the abundance of these rocks in oceanic base-ment. Mantle peridotite exposed at the seafloor and emplaced atcrustal levels may not be true oceanic crust, but in the following weuse the term crust to denote both mafic oceanic crust and peridotiteemplaced at crustal levels in oceanic basement (i.e., less than ~7 kmsubseafloor) to distinguish this material from lithospheric mantle,which extends to ~100 km.

Dick (1989) estimated that (serpentinized) peridotite constitutesat least 13% of exposed basement along the slow-spreading SW Indi-an and Antarctic Ridges. Much has been learned subsequently aboutthe crustal structure at slow- and ultra-slow spreading ridges, butthis estimate remains reasonable. Cannat et al. (1995a) are commonlycited for an estimate of ~25% peridotite in slow-spread oceanic crust,

but they make no such estimate. Instead, these authors show thatareas along the MAR that exhibit a positive gravity anomaly consistof variable proportions of mantle peridotite and mafic crust (basaltand gabbro), and that such thinned and “mixed” crust comprise~25% of a 4000-km2 zone about the MARK area on the MAR. If 50%of these residual gravity anomaly areas is serpentinized peridotiteand this is extrapolated to the entire slow spreading ridge system,then ~12% of slow-spreading crust is serpentinite. Approximatelyhalf of oceanic crust produced at slow-spreading centers is affectedby such segment discontinuities, i.e., is tectonically thinned and poten-tially contains exposures of serpentinized peridotites (Cannat et al.,2010; Tucholke and Lin, 1994). Gabbros are heterogeneously distribut-ed in these oceanic core complexes as large (tens to >100 km2) bodieswithin serpentinized peridotites (e.g., Canales et al., 2008; Ildefonse etal., 2007; McCaig et al., 2010), so if half of the basement in core com-plexes is peridotite, then ~25% of the crust formed at slow spreadingrates is serpentinized peridotite.

Serpentinites were cored at ODP Site 920 on the MAR, and seis-mic velocities increase with depth reaching mantle velocities at3–4 km below seafloor, suggesting a decrease in serpentinizationwith depth (Canales et al., 2000). Seismic evidence for faulting atup to 8 km below the MAR suggests that fluids may penetrate andpartly serpentinize peridotite at great depths, with increasingdegrees of serpentinization as peridotite is uplifted and deformedat shallower depths (Andreani et al., 2007). Cannat et al. (2010) es-timate that the 3.5 km thick section at Site 920 is 60–70%serpentinized on average, but point out the limitations of seismicobservations and that gabbros can have seismic properties similarto partly serpentinized peridotite making them difficult to distin-guish. Using seismic velocity profiles through oceanic crust,Carlson (2001) estimated that Atlantic crust could contain up to13% ultramafic material on average.

Serpentinite can locally comprise significant portions of oceanicbasement. Escartín et al. (2008) suggest that more than 50% of theMAR between 12° 30′N and 35°N consists of basement exposed bydetachment faulting. Kelemen et al. (2007) estimate that 60–80% ofthe region around the 15° 20′N fracture zone on the MAR area isserpentinite, and entire segments of ultraslow spreading ridges arecomprised of ~90% peridotite (Cannat et al., 2006, 2010; Dick et al.,2003; Michael et al., 2003). The observations above thus suggestthat 10–30% are reasonable estimates for the amount of serpentinitein oceanic basement formed at slow and ultraslow spreading rates.

Cannat et al. (2010) provide perhaps the best quantified estimatesof the amount of peridotite in oceanic basement. Using seafloor mor-phology, gravity and crustal thickness data, and sampling of the sea-floor, Cannat et al. (2004, 2010) estimate that 20–25% of seafloorformed at slow spreading rates (b40 mm y−1) consists of areaswith frequent ultramafic outcrops. From sampling these areas, theyestimate that these consist of 65–90% ultramafic rocks, and takinginto account variations in crustal thickness they calculate that 9–32% of basement formed at slow spreading ridges is ultramafic rock.Cannat et al. (2010) use 9% as a conservative estimate for slowand ultraslow spreading ridges, as the higher estimates are formagmatically starved areas, such as 15°N MAR and 61–69°E on theSWIR.

Serpentinized peridotite has also been sampled locally in propagat-ing rifts and in fracture zones in basement formed at fast spreadingrates in the Pacific. At Hess Deepmantle peridotites were serpentinizedand exposed during rifting as the Galapagos spreading center propagat-edwestward into crust formed at the EPR (East Pacific Rise; Früh-Greenet al., 1996). In contrast, peridotites have not been sampled wherethe east rift of the EPR propagates northward into Pacific crust at PitoDeep at ~23°S, associated with the Easter microplate (Hekinian et al.,1996; Perk et al., 2007). Serpentinized peridotites occur in the Garretttransform fault at 13° 30′S on the EPR and the Terevaka transform asso-ciated with the Easter microplate at 24°S (Hebert et al., 1983; Bideau et


al., 1991; Hekinian et al., 1996). Contreras-Reyes et al. (2008) suggestthat crustal thinning around the Mocha Fracture Zone at ~38°S nearthe Chile Trench and low seismic velocities across the Clipperton Frac-ture Zone near 10°S on the EPR (van Avendonk et al., 2001) may be as-sociatedwith serpentinization ofmantle peridotites in Pacific basement.

Of the global plate production of 3 km2 y−1, 77% (or 2.3 km2 y−1)occurs at intermediate and fast spreading rates, and a crustal thick-ness of 7 km yields 16.1 km3 y−1 mafic crust formed at spreadingrates >40 mm y−1. Combining this with the 4.3 km3 y−1 crustal pro-duction at slow spreading ridges (6 km thick, Cannat et al., 2010)gives 20.4 km3crust y−1 (Table 2). Using 9–20% serpentinite inthe 4.3 km3 y−1 crust formed at slow spreading ridges yields 0.39–0.86 km3 y−1 serpentinite. Bird (2003) estimates that approximately1% of constructional plate boundaries comprise transform faults, andif half of thismaterial in transform faults in crust formed at intermediateand fast spreading rates is serpentinite, then this would amountto 0.08 km3 y−1 serpentinite (0.5×0.01×16.1 km3 y−1). Thus, we es-timate 0.47–0.94 km3 y−1 serpentinite production, or 2.3–4.6 vol.%serpentinite in global oceanic basement (Table 2).

3.2. Abundance of serpentinite in subducting slabs

In addition to serpentinization at slow-spreading mid-oceanridges, suboceanic mantle beneath crust formed at intermediate and

Table 2Carbon and sulfur budgets for seafloor serpentinites.

Oceanic crustal production Area km2 y−1

Total 3Intermediate and Fast spreading 2.3aSlow Spreading

Volume of serpentinite Vo

Intermediate and fast spreading 0.5Slow spreading 9–Total 2.3

Proportions of low- vs high-temperature serpentinite

Thickness km

Low-temperature 0.05–0.2High-temperature 1–6

Carbon in serpentinite Concentration wt.%

Cenozoic Me

Low-temperature 0.25 1.4High-temperature 0.027 0.03Total

Sulfur in serpentinite Concentration p

Low-temperature 850–2300High-temperature 320–2000Total

Serpentinite from slab bending at subduction zones (for a 10% serpentinized layer)Layer thickness km Element added

3–10 100–200 ppm S3–10 0.024 wt.% Carb

Dehydration of serpentinite (for 10% serpentinized layer)Layer thickness km Element lost

3–10 260 ppm Sulfur

Subduction of serpentinite dehydration products (for 10% serpentinized layer)Layer thickness km Element concen

3–10 650 ppm Sulfur3–10 0.027 wt.% Carb

Calculations use density of 2.6 g cm3 for serpentinite, and 3 km2 y−1 rate for production ana Cannat et al. (2010).b Only 23–29% of this sulfur is from seawater, the remainder is transferred from altered

fast spreading rates may be serpentinized as the lithosphere flexesoutboard of subduction zones in the Pacific (Peacock, 2001; Raneroet al., 2003). Normal faulting is common on the incoming plate, andtensional events occur at depths of 25–30 km, suggesting that faultscould allow penetration of fluids into the oceanic mantle (Peacock,2001; Ranero et al., 2003). Seismic refraction and gravity measure-ments at the north Chile trench suggest that the incoming lithosphericmantle is ~17% serpentinized to a depth of ~20 km below the oceaniccrust (Ranero and Sallares, 2004). Double seismic zones, separatedvertically by 20–40 km, define the top of the downgoing oceaniccrust and the bottom of an underlying layer that has been interpretedto be serpentinized oceanic mantle in northern Japan and elsewhere(Peacock, 2001). In Japan this layer is tens of km thick, and seismictomography suggests 10–24% hydration of harzburgite (Zhang et al.,2004). The lower plane of earthquakes here is consistent with a changefrom forsterite+enstatite+H2O plus somehydratedminerals (serpen-tine) to anhydrous harzburgite (Zhang et al., 2004). The above seismicdata can thus be considered to reflect variably hydrated mafic crustunderlain by a layer of suboceanic mantle that is ~10–20 km thickand ~10–20% serpentinized (1–3 wt.% H2O total).

Alternative explanations for the lower seismic zone and recent seis-mic work, however, suggest a thinner zone of mantle serpentinizationat subduction zones. The lower seismic zone may result from seismicrupture via periodic shear instabilities (Kelemen and Hirth, 2007) and

Thickness km Volume km3 y−1

20.47 16.16 4.3

lume % Serp. Volume km3 y−1

0.0820 0.39–0.86–4.6 0.47–0.94

Volume % Volume km3 y−1

0.8–16.7 0.004–0.1682.3–99.2 0.47–0.77

Uptake×1011 mol C y−1

sozoic Cenozoic Mesozoic

0.12–0.16 0.68–0.923 0.23–0.48 0.29–0.58

0.35–0.64 0.97–1.5

pm Uptake×1011 mol S y−1

0.02–0.050.11–1.400.13–1.46b

Uptake 1011 mol y−1

ulfur 0.07–0.49on 0.53–1.7

Loss 1011 mol y−1

0.19–0.63

tration Flux 1011 mol y−1

0.5–1.6on 0.5–1.7

d subduction of oceanic crust.

gabbros (see Section 4.2).


anisotropy in anhydrous peridotite (Reynard et al., 2010). Faccenda etal. (2012) suggest that the lower zone may result from hydrofracturingby water sourced from unbending of the slab at greater depths, withanhydrous peridotite present between the double seismic zones.Recent seismic work provides estimates of serpentinization due toslab bending ranging from 10 to 15% serpentinite extending ~3 kminto the mantle at the Middle-America Trench (Grevemeyer et al.,2007), to up to 30% serpentinization extending 7–14 km below thecrust at Middle America and Chile Trenches (Contreras-Reyes etal., 2008; van Avendonk et al., 2011). Based on the observations sum-marized above, we use a 3 to 10 km thick layer consisting of 10%serpentinite in the following Sections 4 and 5, where we discussfluxes of carbon and sulfur resulting from serpentinization outboardof subduction zones (Table 2).

The chemistry of serpentinization reactions in these cases may dif-fer from those at mid-ocean ridges because seawater must penetratethrough and react with mafic crust before reaching the underlyingmantle. At Hess Deep in the eastern Pacific, the upper mantle wasserpentinized as seawater penetrated downward during rifting ofthe crust, and low seawater-rock ratios resulted in strongly reducingconditions, formation of a low-sulfur secondary mineral assemblage,and generally little change in bulk rock sulfur contents or isotopecompositions (Alt and Shanks, 1998; Früh-Green et al., 1996). Similarprocesses may occur during serpentinization of suboceanic litho-spheric mantle outboard of subduction zones where fluids must tran-sit through ~7 km of mafic crust to reach the mantle and fluid/rockratios are probably low. Geochemical evidence, however, suggeststhat the chemical effects of serpentinization of the suboceanic mantlemay be sufficient to influence the compositions of arc volcanics. Forexample, enrichments of boron and high δ11B and δ37Cl in the Izuand central American arcs have been attributed to serpentinite dehy-dration reactions in the subducting slab (Barnes et al., 2009b; Rüpkeet al., 2002; Stern et al., 2006; Tonarini et al., 2007).

partial m

elt

Arc Crust

slab crust

slab

300400500

800

1200

400

D

Li re to S

Serpentinite dehydration tochlorite-harzburgiteS: 0.2-0.6C: 0

Chlorite-harzburgite subductionS: 0.5-1.6C: 0.5-1.7

Fig. 5. Schematic diagram summarizing processes related to serpentinite at mid-ocean ridge(C) given in 1011 moles per year (Table 2, see text). Serpentinite bodies within slab crust rsubduction zones via normal faults as slab bends, and mantle wedge is serpentinized by slasubduction metamorphism. Dashed red line indicates stability limit of serpentine (after Ulupward (arrows) to metasomatize the mantle wedge and induce partial melting. Yellow zonbreakdown of amphibole. Diagram illustrates intermediate subduction zone based on NE Ja

4. Uptake of carbon and sulfur in seafloor serpentinites

In order to estimate the proportions of low- and high-temperatureserpentinite we use available constraints from ophiolites and oceandrilling. Low-temperature serpentinization occurs near the seafloor,with the low-temperature ophiolite samples all from depths b30 mand the seafloor drillholes penetrating b156 m. Thus we use depths of50 and 200 m as minimum and maximum extents of low-temperatureserpentinite.Weassume the remainder is high-temperature serpentinite:as a minimum value we use 1 km, the approximate thickness ofthe thrust sheets in the ophiolite, and as a maximumwe use the entireoceanic crustal thickness of 6 km in such areas (Cannat et al., 2010).Thus, 0.8–17% of the serpentinite (or 0.004–0.16 km3 y−1) is low-temperature serpentinite (Table 2).

4.1. Carbon

The mean carbon contents of low-temperature serpentinitesfrom the Jurassic Apennine rocks and Cretaceous Iberian margin arehigh (1.06–1.72 wt.%) compared to those of young low-temperatureserpentinites from the LCHF and from ODP drill holes (0.18–0.32 wt.%;Table 1). Mafic ocean crust shows similar age variations, which reflecthigher atmospheric CO2 levels and greater carbon uptake by oceaniccrust in the Cretaceous and Jurassic compared to the Cenozoic (Alt andTeagle, 1999; Berner and Kothavala, 2001; Gillis and Coogan, 2011).We use carbon contents of 1.4 and 0.25 wt.% to calculate carbon uptakeduring low-temperature serpentinization in theMesozoic and Cenozoic,respectively. The primary mantle concentration of carbon (30 ppm,or 0.003 wt.%; Dasgupta and Hirschman, 2010) is subtracted fromthe mean concentrations in these calculations. For high-temperatureserpentinites, the modern oceanic samples contain greater amounts ofcarbon than do the Jurassic Apennine ophiolite serpentinites (Table 1).The former are all from relatively shallow drillholes on the seafloor,

amphibole

mantle

100200

100

200

800

1200

istance (km)

Dep

th (

km)slab serpentinite

wedge

serpentinite

zardite/chrysotilecrystallization antigorite: 0 C: 0

Serpentiniteat MORS: 0.13-1.46C: 0.35-0.64

Slab flexureS: 0.07-0.49C: 0.5-1.7

s (MOR) and subduction zones discussed in text. Global fluxes of sulfur (S) and carbonesult from processes at MOR and fracture zones. Suboceanic mantle is serpentinized atb-derived fluids. Oceanic chrysotile and lizardite recrystallize to antigorite during earlymer and Trommsdorff, 1995), and fluids derived from serpentinite dehydration movee of slab crust indicates progressive dehydration of hydrated mafic crust until eventualpan (after Alt et al., 2012a). See text for discussion.


and their higher carbon contents likely reflect the influence of late,low-temperature precipitation of carbonate. Thus we use the lowestmean carbon content of the drill cores (0.027 wt.% for Hole 1268A) formodern high-temperature serpentinization, and the mean ophioliteconcentration (0.033 wt.%) in high-temperature serpentinites to calcu-late carbon uptake in the Mesozoic.

Using these values and the amounts of low- and high-temperatureserpentinite given above, we calculate a carbon sink of 0.35–0.64×1011 mol C y−1 for modern serpentinization, and a higherrate of 0.97–1.5×1011 mol C y−1 for the Mesozoic (Table 2, Fig. 5).This is comparable to the carbon sink of 0.8 to 2.2×1011 mol C y−1

in oceanic serpentinites calculated by Schwarzenbach (2011) usingdifferent assumptions. Based on the abundance of carbonate veinsin serpentinite, Bach et al. (2011) estimate 0.1 wt.% carbon inserpentinite around the 15° 20′ Fracture Zone on the MAR. Usingthis value yields a carbon sink of 0.85–1.86×1011 mol C y−1.While 0.1 wt.% carbon may be reasonable for shallow serpentinitenear the seafloor, it is excessive for deeper serpentinite, leadingto overestimation of the carbon sink in serpentinite. The bulk carbontaken up by sepentinites has δ13C values around −9‰ (using δ13Cvalues of−8‰ and−10‰ for low- and high-temperature serpentinitesfrom Table 1). We compare the carbon sink in serpentinite to that inmafic oceanic crust below.

Alt and Teagle (1999) calculate fixation of 1.5–2.4×1012 mol C y−1

in Mesozoic mafic oceanic crust. The Mesozoic serpentinite carbonfluxes calculated above amount to 5.4–9.3% of the carbon uptake byMesozoic mafic crust. Revising Alt and Teagle's (1999) estimate down-ward by using their carbon content of 0.25 wt.% for a 6 M.y. old, 600 mthick volcanic section (instead of their 2 wt.% for >120 M.y. old crust)yields uptake of 0.45–0.73×1012 mol C y−1 in mafic crust, identicalto other estimates for Cenozoic oceanic basement (Gillis and Coogan,2011). The estimated modern carbon sink in serpentinite amountsto 4–13% of these global Cenozoic carbon uptake rates for mafic oce-anic crust. However, we estimate above that serpentinite comprisesonly 2.3–4.6% of oceanic basement. Thus, the uptake of carbon byserpentinite compared to that bymafic crust is comparable to or greaterthan the proportion of serpentinite in oceanic basement, suggestingthat, on average, carbon uptake by serpentinite is equal to or slightlygreater than that by an equivalent volume of mafic oceanic crust. Asinmafic crust, much of the carbon sink is as carbonate in the uppermost200 m of serpentinite basement.

The annual carbon uptake rates for modern ocean crust calcula-ted above equal 6.8–7.5×1010 mol C km-3peridotite y−1, with1.7–2.6×1010 mol C km−3 peridotite y−1 in low-temperatureserpentinite. In comparison, subaerial weathering of peridotite inthe Semail ophiolite in Oman results in a smaller fixation rate for car-bon, 0.62×109 mol C km−3peridotite y−1, in a 15-m-thick weatheringzone (Kelemen and Matter, 2008). One difference between theseestimates is that the Oman fluxes are for uptake of carbonate, whereasmass balance for the δ13CTotal Carbon in Table 1 indicates that approxi-mately half of the oceanic carbon uptake is as carbonate, with the re-mainder as organic carbon.

A rough estimate can be made for the additional uptake of carbonduring serpentinization of suboceanic mantle as the lithospheric slabflexes at subduction zones. Subtracting the primary mantle carboncontent of 30 ppm (Dasgupta and Hirschman, 2010) from theserpentinites having the lowest mean carbon content in Table 1(0.027 wt.% for Hole 1268A), gives an estimated carbon content of0.024 wt.%. For a 3–10 km thick layer of suboceanic mantle that is 10%serpentinized, this would result in uptake of 0.5–1.6×1011 mol C y−1,greater than that estimated for serpentinization at mid-ocean ridges(Table 2, Fig. 5). Using the lowest measured carbon content of anyserpentinite (~0.01 wt.%) yields an estimated sink of 0.13–0.46×1011 mol C y−1, comparable to serpentinization at spreadingridges. Thus, serpentinization at subduction zones can contributeto subduction fluxes of carbon, but amounts to only about 1% of the

total carbon budget for subducting mafic oceanic crust (SeeSection 5.3.2, below).

4.2. Sulfur

The contents of sulfur vary within each sample set and betweendifferent sample sets (Table 1), contributing to uncertainties in esti-mating sulfur budgets for serpentinization. As outlined below, how-ever, there are similarities between different sample sets thatprovide some confidence in making these estimates.

The mean total sulfur contents for low- temperature serpentinitesfrom the Lost City hydrothermal field, from Holes 1272A and 1274 atthe 15° 20′ N Fracture Zone, and from the northern Apennine ophioliteare similar, 860–940 ppm, whereas the mean from the Iberian Marginis greater, 2330 ppm (Table 1). To estimate sulfur budgets we use totalsulfur contents of 850–2300 ppm for low-temperature serpentinization.

Mean total sulfur contents for high temperature serpentinite samplesets vary more widely, from 320 ppm for Hess Deep to 3850 ppm forHole 1268A (near the 15° 20′N Fracture Zone). Serpentinization duringrifting at Hess Deep occurred at low water/rock ratios and was notinfluenced by hydrothermal activity related to gabbro intrusions, sowe take the mean of 320 ppm for Hess Deep as the low estimate fortotal sulfur in high-temperature serpentinites. The mean values for theAtlantis Massif and the MARK area are generally similar (1420–1755)and lower than that for Hole 1268A, so we use a maximum total sulfurcontent of 2000 ppm for high-temperature serpentinization.

Subtracting a depleted mantle sulfur content of 100 ppm fromthese values (Alt and Shanks, 2003), and using the amounts of low-and high-temperature serpentinites as above, we calculate uptake of0.13–1.46×1011 mol S y−1 by serpentinization of oceanic basement(Table 2, Fig. 5). These values are similar to other estimates of the sul-fur sink in oceanic serpentinites of 0.18–0.38×1011 mol S y−1

(Schwarzenbach et al., 2012) and 0.13–1.9×1011 mol S y−1 (Altand Shanks, 2003). Much of the sulfur added to high-temperatureserpentinites is hydrothermally mobilized from gabbros, however(Alt and Shanks, 2003). Seawater sulfate is quantitatively reducedto sulfide in submarine hydrothermal systems (Alt and Shanks,2003), so by mass balance, a serpentinite δ34S of +5‰ amounts to23–29% seawater sulfur with the remainder being igneous sulfur de-rived from gabbros (using Jurassic to modern seawater δ34S=17–21‰and δ34Smantle=0.1‰). This decreases the estimated amount of seawa-ter sulfur uptake in serpentinite to 0.03–0.42×1011 mol S y−1. Thissink amounts to 2–30% of the seawater sulfur sink in mafic oceaniccrust (~1.3×1011 mol S y−1; Alt and Shanks, 2010), comparableto our estimate for the proportion of serpentinite in oceanic crust(9–20%). This indicates that the sink of seawater sulfur duringserpentinization is comparable to that per unit volume of maficcrust, and that incorporation of serpentinite into oceanic crust haslittle effect on the sulfur budget for altered oceanic crust. The sea-water sulfur sink in serpentinite is only a few percent of the riverinesource of sulfate to the oceans (8.8×1011 mol yr−1) and the sedi-mentary pyrite sink (5.5×1011 mol yr−1; Holser et al., 1988).

Additional sulfur could be taken up by serpentinization duringslab flexure at subduction zones. Addition of 100–200 ppm sulfur toa 3–10 km thick layer of suboceanic mantle that is 10% serpentinizedwould result in a further sink of 0.07–0.48×1011 mol S y−1, compa-rable to the sink in serpentinite at mid-ocean ridges.

5. Effects of subduction metamorphism on carbon and sulfurin serpentinites

5.1. Recrystallization of lizardite and crysotile to antigorite

Alt et al. (2012a,b) report sulfur and carbon data for seafloorserpentinites affected by metamorphism during subduction. TheBeigua unit of the Voltri Massif in Liguria (N. Italy) comprises oceanic


basement similar to the northern Apennine ophiolites, and includesperidotites that were intruded by gabbro bodies and that wereserpentinized on the seafloor (Ernst, 1981; Scambelluri et al., 1997).The Voltri Massif was subducted and metamorphosed at temperaturesof ~550 °C and pressures of 2–2.5 GPa during Cretaceous-TertiaryAlpine tectonic convergence (Ernst, 1981; Messiga and Scambelluri,1991; Mottana and Bocchio, 1975). The seafloor lizardite and chrysotileof the serpentinites are recrystallized to antigorite, and minor sec-ondary olivine and Ti–clinohumite result from partial dehydration ofserpentinite (Scambelluri et al., 1997).

The Beigua antigorite–serpentinites average 690 ppm sulfur hav-ing δ34S=9.3‰ (Table 1). These enrichments of sulfur and 34S arelike those in high-temperature serpentinites from the seafloor andfrom northern Apennine ophiolites, and reflect the influence ofhydrothermal fluids that had interacted with gabbroic rocks beneaththe seafloor (Fig. 3; Alt et al., 2012b). Total carbon contents of theserpentinites (230 to 320 ppm) are similar to those of the other high-temperature serpentinites, and δ13CTotal C values (−10.9 to −8.5‰)fall along the same mixing line between organic carbon and seawatercarbon, as do other serpentinites (Fig. 4; Table 1).

Thus, there is no evidence from the Beigua antigorite serpentinitesfor change in sulfur or carbon during recrystallization of seafloorserpentinites to antigorite serpentinites, which is inferred to take placeat temperatures of ~300–380 °C (Deschamps et al., 2011; Kodolanyiand Pettke, 2011), nor during partial dehydration to olivine at tempera-tures of ~550 °C and pressures of 2–2.5 GPa (Alt et al., 2012b). Littlefluid would be generated during recrystallization to antigorite or duringformation of the minor amounts of olivine and Ti–clinohumite, andthe fluid mobile elements Li, As, Sb, Pb, U, Ba, Sr, and Cs are alsoretained during recrystallization of lizardite and chrysotile to antigorite(Deschamps et al., 2011). Loss of a fluid phase upon heating can becaused by dehydroxylation of antigorite due to polysome variations(Padrón-Navarta et al., 2008; Wunder et al., 2001), but the effect ofthis on element mobilization is not known. Oxygen and hydrogen iso-tope measurements of serpentinites from the Erro-Tobbio peridotiteunit of the Voltri Massif, which has a metamorphic history similar tothat of the Beigua unit, are consistent with the Beigua results. TheErro-Tobbio data indicate preservation of isotope signatures of oceanicserpentinization in peridotite that underwent static recrystallizationto high-pressure assemblages, and only local scale fluid flow at lowwater–rock ratios in shear zones, in a closed system (Früh-Green etal., 2001). High-pressure serpentinites from Monviso in the Alps havevariable sulfur contents (below detection up to 2410 ppm), but havehigh δ34S values of 11.7–17.4‰ that are interpreted to retain the influ-ence of seafloor serpentinization (Hattori and Guillot, 2007), consistentwith results from the Beigua antigorite serpentinites. In the presence of afluid phase, however, some B and Sr may be lost during recrystallizationof lizardite to antigorite (Kodolanyi and Pettke, 2011), and fluxing ofsubducted mafic crust with fluids can result in loss of carbon (Gormanet al., 2006), so in other settings carbon and sulfur in subductedserpentinites may be mobile.

5.2. Dehydration of antigorite serpentinites

The Cerro del Almirez ultramafic complex lies within the BeticCordillera in southern Spain and preserves high-pressure break-down of antigorite–serpentinite to chlorite-harzburgite (olivine+orthopyroxene+chlorite; Garrido et al., 2005; Padrón-Navarta et al.,2011; Trommsdorff et al., 1998). The antigorite–serpentinites are simi-lar to those in the Penninic zone of the Alps (including the Voltri Massif)and result from an early stage of prograde Alpine subduction zonemetamorphism overprinting previously hydrated oceanic mantle(Scambelluri et al., 2001; Trommsdorff et al., 1998).

The Almirez serpentinites contain 9–12 wt.% H2O and average910 ppm sulfur, have mean bulk δ18O and δD values of 8.6‰ (±1.4‰)and−54‰, (±5‰), and δ34S=5.0‰ (Table 1; Alt et al., 2012a). These

data are similar to those for high-temperature seafloor serpentinites(Fig. 3; Table 1), and indicate integrated serpentinization temperaturesof ~200 °C on the Tethyan seafloor, with fluids derived fromhigh-temperature interactions of seawater with gabbro and perido-tite (Alt et al., 2012a).

The chlorite–harzburgites have bulk δ18O values of 8.0±0.9‰,similar to those of the serpentinites, but contain about half the water(5.7±1.9 wt.% H2O), and have lower δD values of −77±11‰ (Altet al., 2012a). The chlorite harzburgites on average contain less sulfurthan the serpentinites and have lower δ34S values (650 ppm, and1.2‰; Table 1). Alt et al. (2012a) show that dehydration of serpentiniteresulted in loss of 5 wt.% H2O having δ18O=8–10‰ and δD=−27to −65‰, and that 260 ppm sulfur was lost as sulfate having δ34S=14.5‰. Thus, dehydration of serpentinite during subduction metamor-phism releases water and sulfur, which can inducemelting and contrib-ute to 18O, D, and 34S enrichments and oxidation of the sub-arc mantlewedge (Alt et al., 2012a).

The contents and δ13C of total carbon in the two rock typesoverlap (Fig. 4; Table 1), reflecting mixing between reduced carbonin the rocks (210 ppm, δ13C≈−26‰) and seawater-derived car-bonate (Alt et al., 2012a). There is no evidence for change in carboncontents or isotope composition with dehydration metamorphism.

The data for serpentinized oceanic peridotites in Table 1 indicatethat they can carry isotopically fractionated water, carbon and sulfurinto subduction zones, and can contribute to metasomatism of themantle wedge. The chlorite–harzburgites contain isotopically frac-tionated sulfur, water and carbon that can be recycled deeper intothe mantle where they may be significant for volatile budgets of thedeep Earth. In the following section we summarize estimates of fluxesof these volatiles during subduction of serpentinites.

5.3. Recycling of carbon and sulfur during subduction of serpentinite

Serpentinization at slow and ultra-slow spreading ridges is impor-tant for seawater-crustal chemical exchange, but at present such crustis only being subducted at the Antilles and Scotia arcs, so the roleof these serpentinites in subduction recycling may be more of re-gional than global importance today. Although poorly constrained,serpentinization of basement that formed at fast and intermediatespreading rates appears to be restricted to fracture zones and somepropagating rifts, making this of local importance where such fea-tures are subducted. Much of the Tethyan seafloor formed at slow-and ultraslow spreading ridges in the Jurassic (Piccardo et al., 2004;Principi et al., 2004; Puga et al., 2011) so as Tethyan basement wassubducted in the Cretaceous the role of MOR serpentinites may havebeen more important than today. Because subduction of serpentiniteformed atMOR is apparently small on a global scale today, in the follow-ing we summarize the potential effects of subduction of serpentinitethat forms in the mantle outboard of subduction zones.

5.3.1. SulfurSerpentinization of mantle peridotites at Hess Deep occurred as the

overlying crust was rifting and fractures allowed seawater fluids accessto the mantle. Serpentinization occurred at low seawater/rock ratios(Früh-Green et al., 1996),most likely after fluids had penetrated throughmafic crust into the mantle. Thus Hess Deep may be an analogue forserpentinization of suboceanic mantle where the lithosphere flexes atsubduction zones. The Hess Deep serpentinites contain 120±70 ppmsulfide-sulfur, with the rest present as seawater sulfate added to therocks (Alt and Shanks, 1998). A 3–10 km thick slab that is 10%serpentinized and contains 100-200 ppm sulfur thus results in subduc-tion of 0.07–0.48×1011 mol S y−1 (for 3 km2 y-1 subduction rate;Table 2, Fig. 5). This is significantly lower than the rate of subductionof sulfur in altered mafic oceanic crust (~2×1012 mol S y−1; Alt,1995) or in sediments (~0.4×1012mol S y−1; Canfield, 2004).


The Almirez serpentinites lose an average of ~260 ppm S (δ34S=14.5‰) during transition to chlorite–harzburgites. Although theremay be large uncertainty making a global extrapolation from onesuite of samples, the Almirez serpentinites are similar to otherhigh-temperature seafloor serpentinites, and these are the onlydata available for the behavior of sulfur during serpentinite dehydra-tion reactions. A 3–10 km thick layer of 10% serpentinite losing260 ppm S would result in a flux of 0.19–0.63×1011mol S y−1 tothe mantle wedge. This is significantly less than the annual flux ofsulfur from subduction volcanoes, 2.5–6.8×1011 mol S y−1 (Halmeret al., 2002). The sulfate released, however, could be important in ox-idizing the mantle wedge (Alt et al., 2012a; Kelley and Cottrell, 2012).

Subduction of serpentinite dehydration products can also be impor-tant for recycling of volatiles into the mantle. Subduction of a 3–10 kmthick layer of 10% chlorite–harzburgite would lead to a flux of 0.5–1.6×1011mol S y−1 into themantle (for chlorite–harzburgite containing650 ppm S; Table 2, Fig. 5). The Almirez chlorite–harzburgites andserpentinites have positive δ34S values, compared to themantle value ofaround 0.1‰, whereas low-temperature serpentinites have negativeδ34S values, so subduction of serpentinite could contribute to sulfurisotope heterogeneities in the mantle, which have been attributed torecycling of sedimentary pyrite and sulfate, and to isotope fractionationduring degassing (Chaussidon et al., 1987; Wilson et al., 1996).

5.3.2. CarbonThe total subduction of carbon is 5.4×1012 mol C y−1 (Kerrick

and Connolly, 2001). Including a 3–10 km thick layer of suboceanicmantle that is 10% serpentinized would add 0.53–1.7×1011 mol C y−1

to the subduction flux of carbon (Table 2, Fig. 5; for a serpentinite com-ponent containing 270 ppm total carbon, see above). The δ13C value ofthe high-temperature serpentinites, around −10‰ (Table 1), maybe a reasonable value for this additional subducted serpentinite car-bon. This serpentinite could contribute to carbon subduction, but itscontribution is less than that of sedimentary organic carbon(~0.8×1012 mol y−1, δ13C=−12 to −30‰; Bebout, 1995; Sano andMarty, 1995; Shaw et al., 2003) and significantly less than subductionof carbon in sedimentary carbonate (1.2×1012 mol y−1, δ13C=0‰;Plank and Langmuir, 1998) or in altered mafic oceanic crust (3.4×1012 mol y−1, δ13C=−4.7‰; Alt and Teagle, 1999; Alt et al., 2012a,2012b).

The δ13C values of the high-temperature serpentinites (~−10‰;Table 1) and of the Almirez chlorite harzburgites (−9.6 to −20.2‰;Alt et al., 2012a, 2012b) are lower than that of the mantle (-5‰;Dasgupta and Hirschman, 2010), so recycling of serpentinite and itsdehydration products into the mantle could produce isotope hetero-geneities that would appear to be due to subduction of sedimentaryorganic carbon (δ13C=−12 to −30‰) (Fig. 5).

6. Conclusions

We summarize analyses of carbon and sulfur in serpentinized sea-floor peridotites, and in similar rocks metamorphosed in subductionzones. We use a simplified classification to divide serpentinizationinto high-temperature and low-temperature processes, and showthat both result in uptake of carbon and sulfur into serpentinizedperidotites.

High-temperature serpentinization typically involves heat andmass transfer from gabbro intrusions, leading to addition of ~2000 ppm(up to >1 wt.%) hydrothermal sulfide having δ34S≈+5 to +10‰ toperidotite during serpentinization. Total carbon contents are 0.008–0.603 wt.% having δ13C values of −3‰ to −17.5‰, resulting from mix-tures of organic carbon and seawater-derived carbonate.

Low-temperature serpentinization occurs in unaltered peridotite orsuperimposed on peridotites partly serpentinized at high-temperatures.Microbial reduction of seawater sulfate generally leads to addition of~800–2000 ppm (up to 1.4 wt.%) sulfur having negative δ34S values,

down to −45‰. Reservoir effects occur locally where the system be-comes closed to seawater, resulting in evolution to positive δ34S values,up to +27‰. At the seafloor and along faults, oxidation of low-δ34Ssulfide minerals by circulating cold seawater can lead to loss of sul-fide, and addition of high-δ34S seawater sulfate leads to high SO4/ΣSand high δ34S values for total sulfur. Low-temperature serpentinitescontain 0.006–7.2 wt.% carbon, having δ13C values of −26 to +2.2‰and consisting of mixtures of organic carbon and seawater-derivedcarbonate.

We estimate that serpentinization at mid ocean ridges is a sinkfor 0.35–0.64×1011 mol C y−1 and 0.13–1.46×1011 mol S y−1, com-parable to or slightly greater than the sinks of these elements per unitvolume of mafic oceanic crust. The volume of serpentinite forming inthe subducting plate at subduction zones may be comparable to orgreater than that at MOR and may significantly affect chemical budgetsfor serpentinization.

The contents and isotope compositions of sulfur and carbon areunaffected by early reactions during subduction metamorphism. Thereis no evidence for change in the contents or isotope compositions of sul-fur or carbon during recrystallization of seafloor lizardite and chrysotileto antigorite and early partial dehydration and formation of minoramounts of olivine. Dehydration of antigorite-serpentinites to chlorite-harzburgites, however, results in loss of 5 wt.% water, and 260 ppmsulfur is lost as sulfate having δ34S=14.5‰. The released volatiles caninduce melting and contribute to 34S enrichments and oxidation ofthe sub-arc mantle wedge. There is no evidence for change in carboncontents or isotope composition with dehydration metamorphism.

Serpentinized oceanic peridotites carry isotopically fractionatedwater, carbon and sulfur into subduction zones. Roughly 0.07–0.48×1011 mol S y−1 is subducted in serpentinite, significantly less thansubduction of sulfur in altered mafic oceanic crust or in sediments(~2×1012 and ~0.4×1012mol S y−1, respectively). We estimate sub-duction of 0.5–1.6×1011 mol C y−1 in serpentinites, less than 3% ofthe total mass of subducted carbon, 5.4×1012 mol C y−1. Isotopicallyfractionated sulfur,water and carbon remain in serpentinite dehydrationproducts, however, and can be recycled deeper into the mantle wherethey may be significant for volatile budgets of the deep Earth.

Acknowledgments

We thank two anonymous reviewers and the editors of this vol-ume for helpful comments that improved the manuscript. JA's workon serpentinites was supported by the US National Science Founda-tion (OCE 9416007, OCE 0424558, and EAR 0809000). The SwissNational Science Foundation is acknowledged for many years ofsupport to GFG and SMB. CG's work was funded by the Spanish“Ministerio de Economía y Competitividad” (grants CGL2009-12518/BTE and CGL2010-14848/BTE), “Junta de Andalucía” (research groupsRNM-145, RNM-131 and grant 2009RNM-4495) and InternationalLithosphere Program (CC4-MEDYNA). CM's and JAPN's research hasbeen supported by EU-FP7-funded Marie Curie postdoctoral grantsunder contract agreements PERG08-GA-2010-276867 and PIOF-GA-2010-273017.

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The role of serpentinites in cycling of carbon and sulfur: Seafloor serpentinization and subduction...

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