Lithos 177 (2013) 226–244

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Sources and cycling of carbon in continental, serpentinite-hosted alkalinesprings in the Voltri Massif, Italy

Esther M. Schwarzenbach a,⁎, Susan Q. Lang a, Gretchen L. Früh-Green a, Marvin D. Lilley b,Stefano M. Bernasconi c, Sabine Méhay a,1

a Institute of Geochemistry and Petrology, ETH Zurich, CH-8092 Zurich, Switzerlandb School of Oceanography, University of Washington, Seattle, WA, USAc Geological Institute, ETH Zurich, CH-8092 Zurich, Switzerland

⁎ Corresponding author at: Virginia TechGeosciences, 4024061, USA. Tel.: +1 540 231 8521.

E-mail address: esther11@vt.edu (E.M. Schwarzenbac1 Now at: Schlumberger, Reservoir Sampling and Analy

0024-4937/$ – see front matter © 2013 Elsevier B.V. All rihttp://dx.doi.org/10.1016/j.lithos.2013.07.009

a b s t r a c t

a r t i c l e i n f o

Article history:Received 13 February 2013Accepted 10 July 2013Available online 17 July 2013

Keywords:Alkaline springsVoltri MassifCarbon sequestrationCarbonate depositsSerpentinization

Active serpentinization systems are abundant onpresent-day Earth and are of increasing interest becausewater–rock reactions lead to alkaline, Ca–OH fluids that have the potential to sequester CO2 and form reduced chemicalspecies (e.g. CH4 and H2) that can support chemolithoautotrophy. We present a study of alkaline (pH 10–12)springs in the Voltri Massif (Italy) that focuses on the sources and cycling of inorganic carbon in the basementrocks, the interacting fluids, and the resulting surface carbonate deposits. Most springs are located in mantlerocks that underwent varying degrees of ocean-floor serpentinization and incorporation of carbon (as carbonateand organic carbon) in the Jurassic, which is preserved during subsequent subduction and uplift onto thecontinent.The springs are fed bymeteoricwater that evolves into alkaline, Ca-rich springwaters that have 2–3 times higherCa concentrations than the adjacent rivers. Our study is consistent with previous reaction path modeling andidentifies the formation of clayminerals in serpentinites that have been altered by alkaline fluids. These stronglyaltered basement rocks contain up to 2 wt.% C in the shallow subsurface, also documented by the presence oflate calcite veins. Concentrations of dissolved inorganic carbon (DIC) of the alkaline fluids are generally low(b16 μmol/L) and δ13CDIC is between −24.2‰ and +1.3‰. We argue that the concentrations and isotopiccomposition of the DIC in the alkaline waters provide evidence for 1) the precipitation of calcium carbonateunder closed-system conditions with respect to atmospheric CO2 and 2) removal of DIC by microbial activityin the subsurface. Late-stage uptake of atmospheric CO2 in the shallow subsurface or at the exit sites subsequenttowater–rock–microbe interactions in the basement result influidswith lower pH and Ca, and enrichment inDICand 13C. At the surface, interaction of the high pH, Ca–OHfluidswith atmospheric CO2 causes precipitation of car-bonates as travertines or crusts on the basement rocks, storing CO2 as calcium carbonate. Carbonate precipitationat the exit sites is strongly dominated by kinetic processes leading to carbon and oxygen isotope compositions aslow as−27.2‰ and−18.7‰, respectively.With increasing distance from the springs, the 13C and 18O content ofthe carbonate increases and the fluid pH changes towards neutral. Our study shows that surficial carbonateprecipitation plays a subordinate role in carbon sequestration, but that the evolution from Mg–HCO3 springwaters to Ca–OH waters removes significant amounts of carbon, presumably in the subsurface. We calculatethat the serpentinites have the capacity to sequester up to 0.50 to 2.05 × 109 mol carbon per year and concludethat they can take up significantly more CO2 than they currently contain.

© 2013 Elsevier B.V. All rights reserved.

1. Introduction

Ultramafic rocks are commonly exposed on the seafloor and oncontinents (e.g. Bernoulli and Weissert, 1985; Cannat, 1993; Cannatet al., 2010; Kelley et al., 2001; Lagabrielle and Lemoine, 1997). Interac-tion with water results in large chemical changes and in fluids that are

44DerringHall, Blacksburg, VA

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high in pH, Ca2+, H2 and very low in Ctotal and Mg2+ (e.g. Boschettiand Toscani, 2008; Kelley et al., 2001; Marques et al., 2008; Neal andStanger, 1983). These highly reactive systems havemajor consequencesfor lithospheric cooling, carbon sequestration, andmicrobial activity. Re-cent studies have shown that the alteration of ultramafic rocks plays animportant role in the global cycling of carbon (e.g. Alt and Teagle,1999; Alt et al., 2012, 2013; Früh-Green et al., 2004; Schwarzenbachet al., 2013). On land, progressive interaction ofmeteoric waterwith var-iably serpentinized peridotites ultimately results in alkaline, Ca–OH-richwaters that have the potential to store significant amounts of CO2 ascalcium carbonate (e.g. Barnes and O'Neil, 1969; Barnes et al., 1967,

227E.M. Schwarzenbach et al. / Lithos 177 (2013) 226–244

1972, 1978; Boschetti and Toscani, 2008; Bruni et al., 2002; Cipolli et al.,2004; Neal and Stanger, 1983). Interaction of Ca–OH waters with atmo-spheric CO2 causes precipitation of carbonate as travertine deposits orcrusts around the outflow areas of these alkaline springs. In the Samailophiolite (Oman), large travertines and carbonate-cemented alluvialdeposits have formed at least during the last 26,000 years, and recentcalculations suggest that up to 105 tons of CO2 is fixed annually ascarbonate in these rocks (Kelemen and Matter, 2008). Thus, with thesocietal need to limit CO2 concentrations in the atmosphere there is anincreasing interest in understanding serpentinization processes and thecarbonation of serpentinite (Andreani et al., 2008; Boschi et al., 2009;Cipolli et al., 2004; Kelemen and Matter, 2008; Klein and Garrido,2011; Matter and Kelemen, 2009). Additionally, the hydrogen andmethane that result from serpentinization provide abundant metabolicenergy (McCollom, 2007), which has led to an increased interest instudying the ability of such environments to host diverse ecosystems(Brazelton et al., 2006; DeChaine et al., 2006; Lartaud et al., 2011;Menez et al., 2012; Ohara et al., 2012; Schrenk et al., 2004). Becausedissolved inorganic carbon concentrations are extremely low in alkalineserpentinization systems, the availability of oxidized carbon may be alimiting factor for both abiogenic hydrocarbon synthesis and for thehabitability of serpentinization environments for life.

On the ocean floor, interaction of seawater with ultramafic rocks re-sults in the incorporation of marine carbonate into the shallow oceaniclithosphere (Früh-Green et al., 2003; Kelley et al., 2001, 2005; Ludwiget al., 2006; Schwarzenbach et al., 2013). Through subduction of thelithosphere, carbon is transported into the mantle. In some locations,parts of the subducted plates are uplifted and incorporated into moun-tain chains, such as in the Alps and the Apennines (Bernoulli andWeissert, 1985; Früh-Green et al., 1990; Piccardo et al., 2004). Ultramaficrocks exposed on continents thus also give the opportunity to study theflux of carbon into the mantle and the preservation of carbon duringsubduction.

The present study is part of a multidisciplinary investigation oftwelve alkaline springs in the Voltri Massif near Genoa, Italy, which in-cludes microbiological, volatile, geochemical and petrological analyses.In the Voltri Massif, highly alkaline springs issue from almost complete-ly serpentinized peridotites to rare lherzolites (Bruni et al., 2002; Chiesaet al., 1975). Where the spring waters exit the rock, carbonates precip-itate forming thick carbonate crusts. Bruni et al. (2002) performedreaction-path modeling and concluded that the Ca–OH waters formby dissolution of serpentinite and concomitant carbonate precipitationin a closed system with respect to CO2. In this manuscript, we focuson the carbon geochemistry of the fluids, basement rocks, and carbon-ate deposits that form around the outflow areas of the springs. The con-ditions around the springs may support a unique microbial communitythat relies on the availability of H2. Thus the goal of this study was todetermine the source and fate of inorganic carbon in all portions ofa serpentinization system including the mantle rocks, the fluids andultimately the carbonate deposits.

2. Geological setting

The studied springs are located within the Gruppo di Voltri west ofGenoa, Italy. The Voltri Massif structurally belongs to the Alpine orogen(Bezzi and Piccardo, 1971; Chiesa et al., 1975; Scambelluri et al., 1991).The Voltri Group can be divided into three lithologic and tectonic units(Fig. 1; Strating, 1991 and references therein):

1. Beigua Unit: Antigorite–serpentinites and eclogitic metagabbros;2. Voltri–Rossiglione Unit: Mesozoic high-pressure calcschists and

metavolcanic rocks;3. Erro–Tobbio Unit: Partly recrystallized antigorite-bearing

metaperidotiteswith km-scale bodies of unalteredmantle rock (partlyserpentinized clinopyroxene-poor lherzolites and harzburgites (e.g.Borghini et al., 2007; Rampone et al., 2005; Scambelluri et al., 1991))

The antigorite–serpentinites of the Voltri Massif formed in thePiemont–Ligurian Ocean by seawater exposure of mantle rocks duringthe Middle to Late Jurassic (Scambelluri et al., 1995; Strating et al.,1993). During subsequent subduction, these rocks reached eclogitefacies at P-T conditions of 2–2.5 GPa and 550–600 °C (Messiga et al.,1995; Scambelluri et al., 1995; Vignaroli et al., 2010 and referencestherein). We sampled high pH, Ca–OH springs at ten sites within theantigorite–serpentinites of the Beigua Unit, at one site (GOR34, withthree exit sites) in the lherzolites of the Erro–Tobbio Unit, and one site(V18) in metabasaltic rocks of the Voltri–Rossiglione Unit (Fig. 1).

2.1. Description of the carbonate deposits

The strongly alkaline and Ca-rich springs form carbonate crusts ofvarious thickness and locally form travertine deposits. Table 1 summa-rizes the occurrence of carbonate deposits at the sampled springs anda selection of the deposits is described in more detail here.

Most of the springs issue from metal pipes 1–5 cm in diameter(Fig. 2A and B), with the exception of GOR34 where the alkaline watersform travertine deposits with several pools (Fig. 2C). Where the springwaters issue from pipes, carbonate crusts of up to several centimetersthickness form underneath the spring outflow (e.g. LER21 Fig. 2Aand BR2 Fig. 2B). The carbonates generally form either as multiple,thin fine-grained carbonate layers (Fig. 2D) or as homogenous, fine-grained carbonate cement (e.g. at spring BR2). Springs C11, PIO14,BR3 and L43 have only precipitated a very small amount of carbonate.At springs C11 and PIO14 this is likely related to the lower pH, whereasspring L43 issues directly into a small pond and BR3 exits onto forestsoil.

At spring BR1, the alkaline water flows from the pipe over the base-ment and through several small pools before reaching the river, forminga honeycomb structure on the surrounding rocks (Fig. 2E). Thisstructure is typically found where slow-flowing alkaline water causescarbonate precipitation on the basement rocks, forming gray to whitefine-grained layers with brown–orange to gray–white surfaces. XRDanalyses show that calcite is the dominant mineral phase in the de-posits, with traces of aragonite and Ca–Mg carbonates (Table 2). Abrownish color of the surface is likely derived from oxidation of Fe. Inthe vicinity of spring BR1, alkaline water trickles over a ~15 m highwall forming 4–10 cm thick carbonate crusts. These deposits (Fig. 2G)also have a corrugated surface and a characteristic layering andare sometimes covered by microbial matter. Importantly, whilevisiting spring BR1 in different seasons and comparing the carbonatedeposits, we observed that carbonate is periodically dissolved,probably when the river floods over the carbonate deposits (see Fig.2K and L).

Spring LER20 flows at around 4 ml/s and forms a small pool under-neath the pipe; carbonate precipitates along the flow path of the alka-line water until it enters the river. These carbonate deposits show thetypical features as found at most locations: a honeycomb structure onthe surface, a typical layering, a brown–orange color, and plantmaterialincorporated into the carbonate crusts (Fig. 2F). At several locationsalong the Leone river, where springs LER18new, LER20 and LER21 arelocated, alkaline water seeps from fractures in the basement leadingto the precipitation of mm to cm thick carbonate crusts (Fig. 2J). Similarto our observations at BR1, in November 2010 carbonate deposits wereless extensive, possibly due to flooding of the deposits resulting in par-tial dissolution of the carbonates in the riverbed (Fig. 2K and L). Alongthe riverbed of the Leone river, calcite veins of up to a few centimeterswidth occur within the antigorite–serpentinite basement. Thesecalcite veins often contain serpentinite fragments indicating that theyare older and were formed during deformation and fracturing of theVoltri Massif.

Spring V18 emerges from an artificial well, which contains nocarbonate deposits. However, some carbonates form at the side of the

Voltri Massif

Serpentinites andmetagabbros (Beigua Unit)

Calcschists and meta-volcanics (Voltri-Rossiglione Unit)

Erro-Tobbio Peridotite unit

Sestri-Voltaggio Zone

Voltri Group

Mt Figogna Unit

Gazzo-Isoverde Unit

Cravasco-Voltaggio Unit

Flysch Units

8°30’

44°30’

other units

Tertiary deposits

Valosio massif

Valosio

Rossiglione

Voltaggio

Voltri

GenovaArenzano

Sassello

Ligurian Sea0 5 km

GOR34PIO14

S70

LER20

LER21LER18new

C11L43

BR1

BR2

BR3V18

N

Fig. 1. Simplified geological map of the Voltri Massif near Genoa, Italy. The locations of the sampled springs in the units of the Voltri Group are indicated by the gray stars. After Federicoet al. (2004) and Borghini et al. (2007).

228 E.M. Schwarzenbach et al. / Lithos 177 (2013) 226–244

river and on rocks located within the river, where alkaline water seepsfrom fractures in the bedrock.

At spring GOR34, outflow of alkaline water from a shear zone inthe lherzolites has formed travertine deposits containing numeroussmaller pools (b1.5 m in diameter; Fig. 2C). The surface of the carbonatedeposits is also characterized by a honeycomb structure and an orange–brown color. Locally a larger honeycomb texture forms along the edgeof larger pools (Fig. 2H). The carbonate deposits also contain layers ofdifferent color and grain size. XRD analyses suggest the dominance ofcalcite, and traces of lizardite and brucite (Table 2).

3. Sampling and methods

3.1. Sample collection

Twelve different springswere studied during fourfield campaigns inMarch andMay 2009, November 2010, andOctober 2011. The siteswereselected on the basis of their high CH4 contents, which are reported inCipolli et al. (2004) as part of an extensive study of the spring watersin the Voltri Massif (Marini and Ottonello, 2002). Here, we use thenames of the springs in accordance with these previous studies andwhich reflect their locations near rivers flowing through the massif(Bruni et al., 2002; Cipolli et al., 2004; Marini and Ottonello, 2002).Table 1 describes the individual springs and the locations of the rocksample collection.

3.1.1. Water samplingSince most of the springs issue from metallic pipes, water samples

were taken with a teflon tube which was inserted as far as possibleinto the pipe, sealedwith parafilm, and subsequently flushed for severalminutes to minimize contact with air before sampling. The exits of thesprings within the travertine pools of GOR34 are visible due to risinggas bubbles from chimney-like deposits, orifices or fractures at the bot-tom of the pools. Sampling was performed in three pools by inserting ateflon tube directly into the spring outflow. Temperature and pH were

measured in the field. 100 ml of water was collected in plastic bottlesfor major element analyses. To measure the isotopic composition andthe concentration of the dissolved inorganic carbon (DIC), 1–8 ml ofwater was sampled with plastic syringes, filtered with a 0.2 μm filter,and injected into Exetainers® that had previously been prepared with~100 μL of phosphorous acid and flushed with helium.

3.1.2. Carbonate samplingCarbonate deposits were collected at several springs and were

drilled with 1–3 mm drills to obtain depth profiles of the crusts. Atsprings BR1, GOR34, and LER20, samples were collected along theflow-path to determine the relation between distance from the springoutflow and isotope composition. At LER20 water was also sampledalong the flow-path to investigate how the chemistry evolves duringexposure to air and precipitation of carbonate.

3.1.3. Sampling of basement rocksAt selected springs, we collected a number of basement rocks that

were in direct contact with the alkaline water (here called Type 2serpentinites andmetabasalts) to compare themwith others not affectedby the spring waters at the time of sampling (here called Type 1serpentinites) and represent the basement through which the springwaters circulate. At site BR1, a drill core was taken on a wall where alka-line water runs over a ~15 m high serpentinite outcrop covered by thincarbonate deposit. Here, a 15 cm long samplewas drilled to examine theeffect of fluid–rock interaction with depth into the rock (Fig. 3).

3.2. Analytical methods

3.2.1. Water analysesMajor element concentrations of the spring waters (Table 3) were

analyzed on a DX-120 ion chromatograph. An IonPac As14 and an IonPacCs12A (4 × 250 mm) were used as separating columns for anionsand cations, respectively (detection limit 0.1 ppm). The δ18O of H2Owas measured on a Thermo Fisher Delta V Plus Isotope Ratio Mass

Table 1Description of the sampled springs and the locations of the rock sampling.

Name ofthe spring

Coordinates (UTM) Tectonic unit Description of the spring Flow rate(in ml/s)

Carbonate sampling Basement sampling

X (m) Y (m)

BR1 482,387 4,921,360 Beigua unit Spring BR1 is located next to a river (diameter metallic pipe = 1–2 cm);travertines form directly underneath the pipe and continue along the waterstream, forming pools until the water reaches the river. 30 m up theriver alkaline waters have formed a thick (~3–6 cm) carbonate crustcovering a 20 m high serpentinite outcrop.

12 Carbonates were taken around the outflowchannel of the spring, at the wall 30 m fromthe spring, and in one of the pools formingunderneath the spring.

Serpentinite was sampled at the outcrop 30 mfrom the spring, where also a 15 cm longdrill core was taken (see Fig. 3).

BR2 482,682 4,922,007 Beigua unit Spring BR2 issues from a metallic pipe (diameter 4–5 cm, Fig. 2B);carbonates are mainly formed directly underneath the pipe, while thewater is collected in a relatively large pond (4–5 m in diameter); a fewcarbonates are formed along the edge of the pond.

492 Carbonates were sampled from underneaththe outflow channel of the spring and fromthe other side of the big pond.

BR3 482,471 4,921,631 Beigua unit Spring BR3 is located in a small forest (diameter metallic pipe = 1–2 cm);no carbonates form at this spring.

19

C11 475,765 4,924,021 Beigua unit At spring C11 sampling is only possible from an app. 25 m long pipe(diameter metallic pipe = 3.5 cm); white filaments grow inside thepipe and around the outflow area of the spring; no carbonatesform at the end of the pipe.

108

GOR34 485,287 4,920,827 Erro–Tobbiounit

At GOR 34 the alkaline springs form travertines with severallarge pools (diameters up to 1.5 m), where the water iscollected (Fig. 2C & H); three individual pools have been sampled.

Carbonates were sampled inside the pools, atthe walls of the pools, and where the waterflows along the travertine into the next pools.

One serpentinized lherzolite was sampledslightly above the travertine within the faultzone that caused extensive serpentine veining.

L43 482,776 4,938,143 Beigua unit Spring L43 is located in a village next to a chapel. There are nocarbonate deposits or outcrops of basement rocks at this spring(diameter metallic pipe = 1–2 cm).

80

LER18new 472,628 4,918,999 Beigua unit Spring LER18new is located within serpentinite at the side of a river(diameter metallic pipe = 1–2 cm); carbonates form on thesurrounding rocks.

Carbonates were sampled from underneaththe pipe outflow and from deposits on thesurrounding rocks.

LER20 472,658 4,918,751 Beigua unit Spring LER20 is located within serpentinite and issues from a metallicpipe (diameter app. 1 cm), in direct vicinity of a forest; carbonateform on the surrounding rock (Fig. 8).

4 Carbonates were taken from directly underneaththe pipe outflow and along a profile. Detailedsampling is described in Fig. 8A.

One serpentinite sample was taken next to thespring, where the basement is unaffected bythe alkaline water.

LER21 472,825 4,919,060 Beigua unit Spring LER21 is located within serpentinite and issues from a metallic pipe(diameter app. 1–2 cm; Fig. 2A). The outcrop directly at the spring is aconclomerate consisting of large serpentinite and very rare maficcomponents. Thick carbonate crusts form along an approx. 3 m long outcrop.

Carbonates were taken from directly underneaththe pipe outflow and from next to the spring.

One metabasic conglomerate sample was takenfrom directly underneath the spring outflow,showing intense w–r interaction, and aserpentinite from the basement opposite the spring.

PIO14 478,523 4,938,256 Beigua unit Spring PIO14 is located at the side of a river within serpentinite and issuesfrom a metallic pipe (diameter = 1.7 cm). White filaments grow insidethe pipe and in the outflow area of the spring water. No carbonates areformed at this spring.

Serpentinite basement was sampled fromslightly underneath the spring, where springwater trickles over the basement.

S70 472,858 4,931,541 Beigua unit Spring S70 is located next to a river within serpentinite and issues froma metallic pipe (app. 1.5 cm diameter). Some carbonate is depositeddirectly underneath the pipe outflow and further along the waterstream of the spring.

Some carbonate was sampled from underneaththe pipe outflow. Thin crusts form aroundfractures where alkalinewater seeps from the rock.

Three serpentinite samples have been taken:one sample from 10 m above the spring and twosamples from next to the spring, where alkalinewater seeps are observed and limited interactionwith some alkaline water could take place.

V18 485,113 4,921,590 Voltri–Rossigli-one unit

Spring V18 is located next to a river, where the spring water iscollected in a very small pool. The direct outflow is not visible.Carbonates form along the river.

Some carbonate was sampled from the sideof the river.

Two altered metabasic rocks were sampled,which are covered by carbonate deposits.

229E.M

.Schwarzenbach

etal./Lithos177

(2013)226

–244

230 E.M. Schwarzenbach et al. / Lithos 177 (2013) 226–244

Spectrometer (IRMS) with an analytical reproducibility of 1σ b 0.07‰.This system is calibrated with the international standards SMOW,SLAP, and GISP, and δ18OH2O is reported in the conventionalδ-notation with respect to VSMOW. In 2009 and 2010, the δ13C of DICwas analyzed on the same Thermo Fisher GasBench II by measurementof the headspace of the vials containing the acidified sample. In 2011,the δ13CDIC was determined by injecting aliquots of the headspaceinto a Gas Chromatograph (GC) equipped with a CO-PoraPlot Q(27.5 mm × 0.32 mm, 10 μm film thickness, held at 100 °C, heliumcarrier gas at 2.5 mL/min) interfaced through a Conflo IV to a ThermoFisher Delta V Plus IRMS. Analytical error of δ13CDIC is calculatedfrom replicate standard measurements as 1σ b 0.3‰. DIC concentra-tions were determined by injecting aliquots of the headspace with acalibrated gas-tight syringe into the same GC instrument, equippedwith a photoionization pulsed discharge detector (PDD). Amounts ofinjected CO2 were determined by comparison with a 7-point dilutioncurve using a standard gas.

Fig. 2. Examples of occurrence and morphology of carbonate deposits at the sampled springsprecipitate underneath the outflow of the pipe. B) Alkaline spring at BR2 (pipe diameter = 4–5is caught. The largest travertine pool has a diameter of ~1.5 m. D) Layering of a carbonate crustover the serpentinite basement. F) Example of carbonate deposits fromBR1 containing high amouon a 10–15 m high wall ~30 m from spring BR1. Corrugations and a brown–orange to gray colorGOR34, forming a larger honeycomb structure. J) Along the river of spring LER20 and LER21, carbofrom fractures in the rocks and initiates carbonate precipitation upon contact with the atmosphe2009 (K) and November 2010 (L) showing that carbonates are periodically dissolved possibly du

3.2.2. Solid carbon analysesTotal carbon (TC) and total inorganic carbon (TIC) contents, as well

as δ13C of the TC, and δ13C and δ18O of the TIC were analyzed on thebasement rocks and the carbonates. TC and TIC contentsweremeasuredon a CM 5012 CO2 coulometer. For TC analyses, the samples werecombusted at 950 °C in an oxygen atmosphere, for TIC analyses, thesamples were reacted with 2 M perchloric acid in a CM 5130 Acidifica-tion Module and in both cases the produced CO2 was analyzed on thecoulometer. TIC contents are calculated assuming that the carbonatewas only present either as calcite or as aragonite. Reproducibility forTC and TIC is better than 1% and 2.5%, respectively, for standards withN1 wt.% C and maximum error for samples with b600 ppm C is ±30 ppm and ±50 ppm, respectively. The non-carbonate carbon (NCC)contents were calculated as the difference between measured TC andTIC contents.

Carbon isotope measurements of total carbon (δ13CTC) and non-carbonate carbon (δ13CNCC) of the basement rocks were carried out on

. A) Spring LER21 issues from a metallic pipe (pipe diameter = 1–2 cm) and carbonatescm). C) Travertine deposits at spring GOR34 form several pools in which the alkaline waterfrom spring BR2. E) Honeycomb structure at spring BR1 formed when alkaline water flowsnts of leaves and other plantmaterial typically from the nearby forests. G) Carbonate depositscharacterize the surface of the carbonates on the wall. H) Edge of one of the pools at springnate deposits occur as brownish crusts on the basement rocks, as alkaline springwater issuesric CO2. K and L) Comparison of extent of carbonate deposits at spring LER 20 between Mayring flooding of the deposits due to extensive rainfall.

Fig. 2 (continued).

231E.M. Schwarzenbach et al. / Lithos 177 (2013) 226–244

a Thermo Scientific Flash Elemental Analyzer (EA) (1112 Series)interfaced with a Conflo IV to a Delta V IRMS. For δ13CNCC analysis,bulk rock samples were reacted with 3 N HCl to remove all acid solublecarbon,washedwith distilled H2O, dried at 60 °C overnight and homog-enized in an agate mortar. Reproducibility of δ13CTC and δ13CNCC isb0.25‰ (1σ). δ13C and δ18O of the TIC were measured on either aThermo Scientific Kiel IV carbonate device or the Thermo FisherGasBench II. Reproducibility is b0.05‰ (1σ) for δ13CTIC, and b0.07‰(1σ) for δ18OTIC. Carbon and oxygen isotope values are reported in thestandard δ-notation relative to the Vienna-Pee Dee Belemnite (V-PDB)standard, except when noted otherwise.

3.2.3. Petrographic analysesThe mineralogy and petrology of the basement rocks was studied

on thin sections by transmitted and reflected light and by electronmicroprobe (EMP) analyses. EMP analyses were carried out on a JEOLJXA-8200 Electron ProbeMicroanalyzer at 15 kV accelerating potential,20 nA current and 1 μm beam size using natural and synthetic mineralstandards. XRD-analyses were conducted on both basement rock andcarbonate powders on a Bruker, AXS D8 Advance. Due to the low

abundance of clay minerals in Type 2 serpentinites, clay spectra couldnot be separated from the bulk rock spectra.

4. Results

4.1. Basement rocks

4.1.1. MineralogyThe mineralogical description and chemical data of the basement

rocks is reported in detail in Schwarzenbach (2011) and is only summa-rized here (Table 4). Most of the springs are hosted by serpentinites,except GOR 34, which is in lherzolites of the Erro–Tobbio Unit, V18 isinmetabasic rocks of the Voltri–Rossiglione Unit, and LER 21 is in a con-glomerate containing serpentinites and metabasic components. Herewe refer to samples that are largely unaffected by the alkaline springsas Type 1, while samples collected in direct contact with alkaline springwater are called Type 2 (Table 4). Detailed locations of the sampledbasement rocks in relation to the springs are described in Table 1.

Type 1 serpentinites are composed predominantly of antigorite(90–98%), which locally preserve the outlines of bastites that replacepyroxene, and typically show a foliation and variable degree of

Table 2Carbonate deposits and travertines: carbon contents, carbon and oxygen isotope composition, and mineralogical descriptions.

Rock sample Sample name Layera Description Comment XRD TC(%C)

TIC(%C)

NCC(%C) b

Carbonate(in wt.%)

δ13CTC(‰)

δ13CTIC(‰)

δ18OTIC

(‰)

Carbonate layers of drilled samplesBR1_1 BR1_1_L1c Serpentinite 20 cm next to the tube cc ± tlc ± atg ± liz 0.77 0.70 0.08 5.8 −16.5BR1_1 BR1_1_L2c Surface cc ± liz 14.05 10.26 3.79 85.5 −25.3 −27.2 −17.8BR1_1 BR1_1_L3c Surface cc ± liz 12.56 11.18 1.39 93.1 −26.0 −26.7 −17.6BR1_2 BR1_2_S1 1 Surface From the wall Ca–Mg-carb 12.12 11.68 0.44 97.3 −23.3 −23.1 −16.0BR1_2 BR1_2_L1 2 cc 12.16 11.64 0.52 97.0 −23.8 −23.4 −16.1BR1_2 BR1_2_L2 3 cc 11.73 11.73 b l.o.d. 98.0 −24.6 −24.3 −17.8BR1_2 BR1_2_L3 4 cc 11.94 11.81 0.13 98.4 −24.7 −24.4 −17.7BR1_2 BR1_2_L4 5 cc 11.80 11.80 0.00 98.4 −24.4 −24.1 −17.9BR1_4 BR1_4_L4 1 Surface Soaked by water at the edge

of a pool next to rivercc 12.05 11.48 0.57 95.6 −24.4 −23.8 −16.5

BR1_4 BR1_4_L3 2 cc 11.84 11.35 0.49 94.6 −25.8 −25.7 −18.7BR1_4 BR1_4_L2 3 cc 11.92 11.43 0.49 95.2 −25.7 −25.3 −17.9BR1_4 BR1_4_L1 4 cc ± liz ± qz 11.38 10.95 0.43 91.3 −24.7 −24.4 −16.5BR1_D2_S1 BR1_D2_S1_L4 1 Surface From the wall cc 11.85 11.84 0.01 98.7 −20.4 −19.5 −12.2BR1_D2_S1 BR1_D2_S1_L3 2 cc 12.00 11.55 0.46 96.2 −21.8 −21.2 −14.5BR1_D2_S1 BR1_D2_S1_L2 3 cc 11.91 11.56 0.35 96.3 −24.2 −23.3 −16.2BR1_D2_S1 BR1_D2_S1_L1 4 cc 11.82 11.60 0.22 96.6 −23.0 −22.6 −15.9BR2 BR2_L1 1 Underneath surface At the end of the big pond Ca–Mg-carb ± arag 11.94 11.50 0.44 95.8 −15.5 −14.7 −3.7BR2 BR2_L2 2 Mg-cc ± arag 11.48 11.23 0.25 93.6 −14.9 −13.9 −2.8BR2 BR2_L3 3 Mg-cc ± arag 11.53 11.34 0.19 94.5 −14.8 −13.6 −2.9GOR34_1 GOR34_1_L1 1 Surface Wall of travertine Mg-cc 11.75 11.36 0.39 94.7 −25.5 −24.6 −16.3GOR34_1 GOR34_1_L2 2 cc 17.35 11.06 6.29 92.2 −22.1 −21.5 −14.5GOR34_2 GOR34_2_S1 1 Surface Wall of travertine cc 11.89 11.78 0.11 98.2 −20.9 −25.1 −16.6GOR34_2 GOR34_2_L1 2 2 mm below S1 cc 11.12 11.47 b l.o.d. 95.6 −23.4 −20.2 −13.1GOR34_2 GOR34_2_L2 3 1 cm below S1 cc ± liz ± qz 11.42 11.31 0.11 94.2 −25.7 −23.3 −16.6GOR34_3 GOR34_3_L5 1 Surface Soaked in water at inside

edge of the poolcc ± brc ± liz 11.31 11.17 0.14 93.1 −20.5 −19.5 −10.8

GOR34_3 GOR34_3_L4 2 cc ± brc 10.76 10.53 0.23 87.8 −17.0 −15.5 −5.8GOR34_3 GOR34_3_L3 3 cc ± liz ± brc 10.56 10.32 0.24 86.0 −18.7 −17.5 −8.2GOR34_3 GOR34_3_L1 4 cc ± liz 10.80 10.75 0.06 89.5 −20.7 −19.6 −10.7GOR34_3 GOR34_3_L2 5 cc ± liz 11.31 11.19 0.12 93.2 −22.6 −21.9 −12.2GOR34_4 GOR34_4_L3 1 Surface Soaked in water at outer edge

of the poolcc 11.14 11.12 0.02 92.7 −19.8 −18.9 −9.4

GOR34_4 GOR34_4_L2 2 cc ± liz ± brc 10.24 9.93 0.31 82.8 −17.3 −16.2 −7.3GOR34_4 GOR34_4_L1 3 1.3 cm below L2 cc ± liz ± brc 10.68 10.38 0.30 86.5 −16.4 −15.3 −6.3LER18new_1B LER18new_1B_L3c 1 Surface Directly underneath tube cc 13.29 11.01 2.28 91.7 −24.2 −25.2 −16.5

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LER18new_1B LER18new_1B_L2c 2 cc ± liz 12.22 11.19 1.03 93.3 −24.9 −24.9 −17.0LER18new_1B LER18new_1B_L1c 3 cc ± liz ± melilite 11.72 11.00 0.72 91.6 −23.5 −22.4 −17.2LER18new_3 LER18new_3_L4 1 Surface 1.5 m to the right of the spring,

now weakly soaked byspring water

cc 14.14 10.52 3.62 87.7 −25.6 −26.0 −18.2

LER18new_3 LER18new_3_L5 2 Directly below L4 cc 13.23 11.10 2.12 92.5 −25.3 −25.2 −17.2LER18new_3 LER18new_3_L3 3 cc ± liz 11.76 11.17 0.58 93.1 −25.0 −24.2 −16.6LER18new_3 LER18new_3_L2 4 cc 11.82 11.76 0.06 98.0 −23.3 −22.6 −16.5LER18new_3 LER18new_3_L1 5 cc 11.78 11.52 0.26 96.0 −23.6 −22.6 −16.9LER20_2 LER20_2_L5 1 Below surface Soaked by water directly

next to the tubecc 12.35 11.11 1.24 92.6 −20.5 −19.6 −11.1

LER20_2 LER20_2_L4 2 cc 12.14 11.44 0.71 95.3 −21.4 −20.5 −10.5LER20_2 LER20_2_L3 3 cc 12.05 11.58 0.47 96.5 −21.3 −20.4 −11.2LER20_2 LER20_2_L2 4 cc 11.92 11.56 0.36 96.4 −20.8 −20.3 −9.4LER20_2 LER20_2_L1 5 cc 11.99 11.52 0.48 96.0 −21.0 −20.1 −9.8LER21_1 LER21_1_L3c 1 Surface Directly underneath the tube cc 12.20 11.41 0.79 95.1 −26.4 −25.8 −17.4LER21_1 LER21_1_L2c 2 cc 12.07 11.65 0.42 97.0 −27.3 −26.1 −17.7LER21_1 LER21_1_L1c 3 cc 12.09 11.59 0.51 96.5 −25.6 −25.0 −16.8V18_1 V18_1_L4 1 Surface Underneath weaker seapage

close to the springcc 12.46 10.34 2.12 86.2 −21.2 −20.7 −14.7

V18_1 V18_1_L3 2 cc ± atg ± liz ± ol 7.84 6.61 1.23 55.1 −20.3 −18.8 −13.3V18_1 V18_1_L2 3 cc 11.06 9.83 1.24 81.9 −19.4 −18.7 −13.6V18_1 V18_1_L1 4 cc 10.87 9.29 1.58 77.4 −17.8 −16.5 −12.4

Carbonate bulk rock samplesBR1_E BR1_E_1 Top layer, bulk rock Carbonate from 2 m

from spring11.85 10.87 0.97 90.6 −25.0 −23.6 −15.7

BR1_E BR1_E_2 Middle layer, bulk rock 11.76 10.88 0.87 90.7 −24.2 −23.1 −14.3BR1_E BR1_E_3 Inner layer, bulk rock 10.87 9.95 0.92 82.9 −20.6 −20.4 −13.5BR1_5 BR1_5 (pool) Top few cm, bulk rock Sampled 5 m from spring;

barrier of a small poolcc ± liz (traces) 11.32 11.11 0.21 92.6 −22.5 −21.7 −14.1

BR2_A BR2_A Cement from directly underspring outflow

cc ± liz 3.33 3.20 0.13 26.7 −8.3 −9.1 −7.9

BR2_B BR2_B Top layer, bulk rock From under the spring outflow cc ± liz (traces) 10.66 10.34 0.32 86.2 −21.5 −20.8 −13.0BR2_C1 BR2_C1 Top layer, bulk rock From under the spring outflow cc ± liz (traces) 11.20 11.07 0.13 92.2 −24.9 −24.3 −15.8BR2_C2 BR2_C2 Lower layer, bulk rock From under the spring outflow cc ± liz (traces) 11.31 10.97 0.34 91.4 −22.7 −21.1 −13.0LER20_A LER20_A 4 m below the spring cc 11.49 11.49 0.00 95.7 −20.0 −19.6 −9.2LER20_B LER20_B 2.5 m below the spring;

fast fluid flowcc ± liz (traces) 11.64 11.45 0.19 95.4 −18.7 −17.9 −5.9

LER20_C LER20_C 1.2 m below the spring cc 11.69 11.58 0.11 96.5 −19.0 −18.6 −6.7GOR34_2T_C GOR34_2T_C From inside a travertine pool 9.96 9.59 0.38 79.9 −21.3 −20.0 −12.5

a Layers within the profile: 1 = surface, 4 = lowest layer.b NCC calculated from TC and TIC measurements as NCC = TC–TIC.c Carbonate deposits sampled directly underneath the outflow channels of the springs and therefore characterized by instant carbonate precipitation.

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deformation. They contain low amounts (b5%) of olivine and pyrox-ene that show incipient serpentinization along the rim and areovergrown by recrystallized antigorite that either formed duringhigh-P metamorphism or during serpentinization by meteoricwater in the subsurface. Olivine has an XMg of 0.89 to 0.96, whileclinopyroxene has a diopsidic composition (23.8 to 26.9 wt.% Ca).Spinel, magnetite and sulfides (pentlandite, millerite, violarite,heazlewoodite, godlevskite) occur in all samples in low amounts(b5%). Only a few samples show traces of calcite veins cuttingthe groundmass, while in some Type 1 serpentinites in the vicinityof the springs, late fine-grained brown to yellow veins cut the ser-pentine groundmass consisting of very fine-grained clay minerals(saponite to sepiolite) and brucite to Fe-rich brucite (Fig. 4).

Type 2 serpentinites at spring BR1 are characterized by a foliationand show strong alteration of the typical mineral assemblage observedin Type 1 serpentinites. While serpentine still dominates the rocks,chlorite, clay minerals and amphibole replace serpentine as veins andin bastites, respectively. Serpentine defines two generations; a fine-grained groundmass probably formed during seafloor serpentinization,and idiomorphic needles in parts overgrowing amphibole and clayminerals. Due to the very fine-grained texture the clay mineralscould not be determined unequivocally. EMP analyses indicate thepresence of fine-grained mixtures between serpentine + saponiteand serpentine + sepiolite (Fig. 4). We relate the formation of clayminerals to interaction with alkaline water, since clay minerals formedduring seafloor weathering would have been replaced during subse-quent eclogite faciesmetamorphism. Some of the serpentine overgrowsthe clay minerals, therefore, we infer that some of the serpentineformed as a product of continental serpentinization. Late, wide serpen-tine veins cut the groundmass and porphyroblasts, and fine-grainedcalcite veins cut all older structures, which suggests that the calciteveins represent the latest stage of alteration by the alkaline fluids.

Type 2 metabasic rocks at spring V18 are foliated and have afine-grained texture containing biotite + chlorite + plagioclase +epidote + calcite + graphite. Graphite bands and aligned chloritegive the rock a characteristic foliation, while plagioclase formsxenomorphic lenses along the foliation. Calcite occurs as fine-grainedaggregates along the foliation. Similar metabasic rocks are compo-nents of the conglomerate at spring LER 21, which is covered by athick carbonate crust. The sample from LER 21 has a characteristicfoliation with a fine-grained texture and contains the assemblagealbite + K-feldspar + amphibole (actinolite, Mg–hornblende, andFe–hornblende) + Fe–Mg chlorite + epidote + titanite + calcite.Amphibole occurs as needles together with biotite and chloriteelongated along the foliation.

Spring GOR34 is situated within the Erro–Tobbio Unit and ishosted by variably serpentinized lherzolites that contain olivine +pyroxene + spinel + serpentine ± amphibole ± chlorite. The springsare located next to a major fault zone characterized by strong serpentineveining. Primary olivine shows replacement by serpentine along the grainboundaries forming a mesh texture with small grains of magnetite. Fine-grained intergrowths of talc and serpentine or rare amphibole replaceortho- and clinopyroxene. Spinel is partly replaced by fibrous chloriteforming halos around the spinel grains. Some sections of the rock arecharacterized by strong, parallel serpentine veining. These veins aredominated by serpentine but contain variable amounts of chlorite andfine-grained talc. It could not unambiguously be determined whetherthe alteration mineralogy in these rocks formed during interaction withalkaline water or is related to an earlier event.

4.1.2. Carbon contents and isotope compositions of the host rocksThe geochemical data are listed in Table 4. The carbon contents in

the ultramafic rocks are highly variable, with TC = 109–18,390 ppm,TIC = 63–18,114 ppm, and NCC b 990 ppm. The δ13CTC has a widerange of −16.6 to +0.8‰, while the δ13CNCC is generally depleted in13C and has values between −27.4 and −16.6‰. The δ13C of the

inorganic carbon could only be determined for the high TIC samplesand showed an average composition of +0.6 ± 0.6‰. The twometabasalt samples generally have higher TC, TIC and NCC contents,with TC = 7430–20,615 ppm, TIC = 5641–6414 ppm and NCC =1016–14,975 ppm, while the δ13CTC and the δ13CNCC are less depletedin 13C in the metabasic rocks than in the ultramafic rocks(δ13CTC = −5.6 and −1.8‰, δ13CNCC = −10.3 and −6.0‰; Table 4).Type 2 samples have up to 100 times higher TIC contents than Type 1samples (Fig. 5). Sample V18_3 has the highest NCC content. The non-carbonate carbon consists of organic carbon aswell as elemental carbonin the form of graphite. Graphite could only be detected in sampleV18_3, where the δ13CNCC is −10.3‰. However, a δ13CNCC value of−6.0‰ in sample LER21_3_B also indicates the presence of a carboncompound that is more enriched in 13C than common organic carbon,which generally has an isotopic composition between−30 and−20‰.

The 15 cm drill core from spring BR1 (Fig. 3) shows no distinct geo-chemical trend with increasing depth (Table 4) and is dominated byhigh TIC contents of up to 1.81 wt.%. The δ13CTIC increases slightlyfrom the surface inward, from +0.7 to +1.3‰, while the δ13CNCC hasa relatively constant value of −26.9 ± 0.3‰.

4.2. Spring water chemistry

The pH of the springs was between 9.8 and 12.2 (Table 3). The low-est pH of 9.8wasmeasured at site C11, where thewater flows through a25 m long metallic pipe before sampling is possible. Presumably, thewater equilibrates with air as it travels down the pipe, and the invasionof atmospheric CO2 into the fluids results in a lower pH. The tempera-tures of the springs ranged from 12.1 to 22.5 °C. Concentrations ofcalcium are high in all springs (b57.82 mg/L; Table 3). Magnesiumconcentrations are very low (b7.62 mg/L) and correlate negativelywith pH; in many cases Mg was below the detection limit.

Due to very low inorganic carbon contents in the fluid, the δ13CDICcould not be determined for all samples and in each year we sam-pled. DIC concentrations were only determined in 2011 (Table 3).Concentrations ranged from 8.8 to 15.6 μmol C/L with the exceptionof PIO14, where DIC concentrations were N1000 μmol C/L (Fig. 6).These concentrations are very similar to those observed in previousyears by Cipolli et al. (2004). Where measured, δ13CDIC showed awide range from −24.2 to +1.3‰. The most negative values weremeasured at GOR34, BR3, and V18. For the low-DIC fluids, the most13C-enriched δ13CDIC values were observed in the fluids with the low-est DIC concentrations (r2 = 0.58, p = 0.05; Fig. 6). The δ13CDIC

values do not correlate with pH (r2 = 0.17, p = 0.08). Samples ofthe rivers at BR2 and L43 had DIC concentrations of 2122 and930 μmol/L and δ13CDIC values of −15.6 and −11.2‰, respectively(Fig. 6, Table 3).

4.3. Carbonate deposits

The carbonate deposits comprise 80–99% carbonate (Table 2). The or-ganic carbon contents range from 10 ppm to 6.28 wt.% C. Plant leaves orgrass growing around the springs are often trapped in the carbonates andcontribute to the organic carbon contents of almost all the deposits, eventhough efforts were made to avoid visible plant material in the analyses.The δ13C and δ18O of the carbonates vary between −27.2 and −13.6‰,and −18.7 and −2.8‰, respectively (Fig. 7). The δ13CTC is dominatedby inorganic carbon and is usually within 1‰ of the δ13CTIC (Table 2).No systematic trendswere present in the depth-profiles across the crusts,but variations of up to 6.4‰ could be observed.

At LER20, fluid and carbonate samples were collected at increasingdistance from the spring to investigate the role of fluid equilibrationwith atmospheric CO2 (Fig. 8A). The carbonate sample closest to thespring was the most depleted in both 13C and 18O (δ13C = −26.1‰and δ18O = −17.7‰); with distance from the outlet, δ13C increasedto −17.9‰ and δ18O ranged from −11.2 to −5.9‰ (Figs. 8B and 9).

Fig. 3.Near spring BR1 a drill core was taken from awall upstream from the spring, where alkaline water overflows a ~15 m highwall and forming a 4–10 cm thick carbonate crust. Onlythe uppermost 15 cmof the serpentinite basement could be drilled. Samples A, C1, D1 and D3were analyzed for their carbon geochemistry to obtain optimal variationwith depth. SampleC4 was used for microscopic analyses.

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Likewise, the water samples taken with increasing distance from thespring outflow show clear chemical trends; Na decreases while Mgincreases with increasing distance from the spring (Fig. 9). The changesin concentrations are large between the spring water and the firstsampling location, and then become less pronounced. The increase inCa concentrations at ~1 m from the spring could indicate that smallamounts of spring water emerge in the vicinity of that sampling loca-tion, likely through small fractures in the basement rock.

The carbonates and fluids sampled at site BR1 showed a similarpattern (Table 5). Carbonate sample BR1_1 taken underneath thepipe was strongly depleted in 13C and 18O with δ13CTIC = −27.2‰and δ18OTIC = −17.8‰, and increased to δ13CTIC = −21.7‰ andδ18OTIC = −14.1‰ at 5 m distance. Water samples taken at the samelocations also showed a decrease in Ca and Na, and an increase in Mgconcentrations with distance from the spring source.

5. Discussion

5.1. Evolution of the alkaline spring waters

Serpentinization of ultramafic rocks can lead to the formation offluids with high pH (9–12), high Ca and very low Mg concentrationsand generally very low inorganic carbon contents (Boschetti andToscani, 2008; Kelley et al., 2001; Marques et al., 2008; Neal andStanger, 1983). In these systems, the high pH is mainly the result ofthe dissolution of forsterite and enstatite according to Eqs. (1) and (2)(e.g. Palandri and Reed, 2004):

Mg2SiO4 foð Þ þ 2 H2O→2Mg2þ þ SiO2 aqð Þ þ 4 OH− ð1Þ

MgSiO3 enð Þ þ H2O→Mg2þ þ SiO2 aqð Þ þ 2 OH− ð2Þ

Similarly, the dissolution of serpentine also leads to an increase in pH(Eq. (3)):

Mg3Si2O5 OHð Þ4 serpð Þ þ H2O ¼ 3Mg2þ þ 2SiO2 aqð Þ þ 6 OH− ð3Þ

Thus, high-pH fluids can result not only from interaction of meteoricwater with fresh ultramafic rocks such as the Erro–Tobbio lherzolites,but also with the serpentinized peridotites that dominate the basementrocks of the Voltri Massif.

Bruni et al. (2002)modeled the reaction path of springwaters of theVoltri Massif. They suggest that the springwaters progressively developfrom Mg-rich, SO4–Cl waters to slightly evolved, neutral Mg–HCO3

waters to mature high-pH, Ca–OH waters, which have low Mg and thelowest dissolved inorganic carbon (DIC) concentrations. This evolutionis caused by an initial undersaturation of meteoric water with respect

to serpentine and olivine, resulting in dissolution of these mineralsand an increase in Mg-concentrations. At the same time, Ca is releasedfrom diopside or tremolite. Under open-system conditions with respectto CO2, Ca reacts with CO2 to form calcite. Calcium is continuously sup-plied to the systemby dissolution of serpentinite and the concentrationsare controlled by calcite precipitation. During further water–rock inter-action under more closed-system conditions, CO2 becomes increasinglylimited, resulting in a strong increase in Ca-concentrations when thesystem is depleted in C by carbonate precipitation. Concomitantly, thepH increases due to further dissolution of olivine, pyroxene and/or ser-pentine (Eqs. (1), (2), and (3)). The model calculations of Bruni et al.(2002) showed that the Ca–OH waters are oversaturated with respectto serpentine (Mg3Si2O5(OH)4). Precipitation of serpentine from thestrongly evolved waters consequently leads to very low Mg concentra-tions characteristic of the Ca–OH waters.

5.2. Mineralogical and carbon geochemical effects of alkaline fluids on thebasement rocks

5.2.1. Mineral alterationType 2 serpentinites are characterized by the presence of serpentine,

amphibole, chlorite and fine-grained clay minerals. Chemical analysessuggest that the clays are sepiolite to saponite forming a fine-grainedintergrowth with serpentine (Fig. 4). Since any clay minerals formedduring seafloor weathering would have been replaced during eclogitefacies metamorphism we infer that the observed clay minerals formedas a result of interaction with meteoric water either precipitating to-gether with serpentine or replacing it. The formation of clay mineralsand possible formation of serpentine agrees with the simulations byBruni et al. (2002) – the latterwas described above –who in accordancewith our observationsmodeled the precipitation of a sapolite solid mix-ture and sepiolite amongst other minerals. The presence of serpentineneedles overgrowing the groundmass, however, cannot unequivocallybe related to continental serpentinization, since eclogite facies meta-morphism could have resulted in similar textures. In the lherzolites,formation of serpentine, talc and chlorite could be a product of eithercontinental serpentinization or of an earlier event during tectonicemplacement of the Erro–Tobbio Unit. Similarly, the mineralogy of themetabasic rocks cannot clearly be assigned to interaction with alkalinefluids.

5.2.2. Origin of carbon in the basement rocksMultiple processes also impact the carbon content and isotope com-

position of the basement rocks within the Voltri Massif. Exposure of theserpentinites to seawater during the Jurassic resulted in the interactionwith seawater-derived fluids and the incorporation of carbonate andorganic carbon in the serpentinites. Subduction and subsequent uplift

Table 3Major element concentrations and isotopic compositions of the alkaline, Ca–OH springs of the Voltri Massif.

Sample name pH T (°C) Namg/L

Mgmg/L

Camg/L

Caμmol/L(as HCO3

−)

DICμmol/L

δ13C DIC

(‰)VPDB

δ18O H2O

(‰)VSMOW

BR1a January 2001 11.9 12.0 23.7 0.20 47.40 1185BR1a May 2002 11.8 14.0 23.5 0.00 47.30 1183BR1 March 2009 12.0 12.1 23.2 n.d. 41.20 1030BR1 May 2009 11.9 17.8 23.6 n.d. 40.48 1012BR1 November 2010 12.0 14.8 23.5 n.d. 44.25 1106 1.3 −6.8BR1 October 2011 12.3 13.7 24.9 0.24 34.16 854BR2a September 2001 11.7 20.3 41.1 0.00 61.90 1548BR2 March 2009 11.9 20.5 42.0 n.d. 57.82 1445BR2 May 2009 11.8 22.2 41.4 n.d. 53.03 1326BR2 November 2010 12.0 20.3 41.9 0.24 55.41 1385 −7.0BR2 October 2011 12.3 19.0 42.8 n.a. 44.15 1104 9.9 −6.9BR2 river October 2011 7.9 13.5 6.3 24.59 21.84 546 2122.4 −15.6BR3a May 2002 11.7 13.0 18.4 0.00 40.20 1005BR3 March 2009 11.8 13.0 18.5 n.d. 38.12 953 −10.8 −7.0BR3 May 2009 11.8 13.0 18.1 n.d. 37.36 934BR3 November 2010 11.8 13.7 18.9 n.d. 36.26 907 −19.9 −6.6BR3 October 2011 12.2 12.7 19.7 n.a. 29.46 737 12.7 −19.9C11a January 2001 10.5 10.5 12.8 5.80 3.20 80C11 March 2009 10.0 12.5 12.9 6.04 3.74 93 −8.4 −6.4C11 May 2009 9.8 17.1 24.5 7.62 4.06 101 −9.0 −6.6C11 November 2010 10.2 13.6 13.5 6.34 3.52 88 −8.7 −6.5GOR34a October 2001 11.7 18.5 18.5 0.55 59.00 1475GOR34Aa October 2001 11.6 19.0 18.3 1.81 46.90 1173GOR34_1 March 2009 11.8 14.8 n.d. n.d. n.d. −16.5 −8.3GOR34_1 May 2009 11.9 18.4 18.2 n.d. 52.86 1321GOR34_1 November 2010 12.2 17.2 18.4 0.68 44.28 1107 −17.6 −8.1GOR34_1 October 2011 12.2 16.4 19.0 0.37 37.39 935 13.7 −17.8GOR34_2 March 2009 11.6 13.5 17.5 1.95 41.72 1043 −21.6 −7.8GOR34_2 May 2009 11.7 20.7 19.0 0.10 33.83 846GOR34_2 November 2010 12.0 18.4 19.1 0.26 51.75 1294 −15.4 −8.2GOR34_2 October 2011 12.3 17.7 20.2 0.31 33.32 833 15.6 −24.2GOR34_3 May 2009 11.7 18.0 18.6 n.d. 45.00 1125 −14.5 −8.5GOR34_3 October 2011 11.8–12.2b 18.7–20.3b 20.1 n.a. 40.67 1017 8.8 −10.8L43a January 2001 11.5 23.0 28.3 0.01 49.30 1233L43a May 2002 11.6 22.4 27.7 0.00 49.50 1238L43 March 2009 11.8 20.9 28.5 n.d. 43.66 1092L43 May 2009 11.7 22.5 28.2 n.d. 38.78 970L43 November 2010 11.8 21.5 29.0 n.d. 44.21 1105 −7.4L43 October 2011 12.2 20.4 30.3 n.a. 39.13 978 10.1 −5.4L43 river October 2011 9.0 12.1 5.1 10.65 16.49 412 930.1 −11.2LER18new Mai 2009 11.7 15.5 7.2 n.d. 32.82 820LER20a January 2001 11.6 13.7 12.7 0.16 44.10 1103LER20 May 2009 11.7 15.5 13.0 n.d. 32.15 804LER20 November 2010 11.8 16.1 13.5 0.25 33.48 837 −6.9 −7.2LER20_C November 2010 9.4 n.d. 9.4 20.29 34.88 872 −6.5LER20_B November 2010 n.d. n.d. 10.2 16.34 25.14 629 −6.7LER20_A November 2010 n.d. n.d. 10.3 15.36 24.83 621 −6.6LER20 October 2011 12.2 16.2 13.6 0.26 29.18 729 14.2 −11.8LER21a September 2001 11.5 18.0 9.9 0.02 43.50 1088LER21 Mai 2009 11.7 17.5 9.9 n.d. 30.91 773PIO14a October 2001 10.7 14.0 53.0 4.71 0.60 15PIO14a May 2002 10.5 13.1 53.6 4.50 0.60 15PIO14 March 2009 10.7 12.1 53.1 4.98 1.05 26 −3.4 −8.0PIO14 May 2009 10.7 14.4 53.0 4.91 0.87 22 −3.1 −8.3PIO14 October 2011 11.2 12.0 54.7 5.11 1.28 32 1036.4 −4.4S70a January 2001 11.4 12.0 5.5 0.04 36.40 910S70a October 2001 11.5 13.8 5.4 0.04 36.60 915S70 Mai 2009 11.6 13.3 5.7 n.d. 29.75 744V18a January 2001 11.4 14.0 16.6 0.01 22.50 563V18a October 2001 11.3 15.6 16.1 0.00 22.20 555V18 March 2009 11.5 13.2 16.4 0.34 22.29 557 −19.0 −6.6V18 May 2009 11.5 16.3 16.4 n.d. 18.21 455 −19.0 −8.0

a Data from Cipolli et al. (2004).b Range of values during 3-hour sampling period.

236 E.M. Schwarzenbach et al. / Lithos 177 (2013) 226–244

onto the continent eventually led to interaction with meteoric fluidsthat additionally affected the carbon geochemistry.

Type 1 serpentinites have carbon contents and isotope signatures sim-ilar to oceanic serpentinites. Their total carbon content ranges from109 to1380 ppm C, with inorganic carbon contributing between 45 and 82% of

the total carbon (average 58% ± 12%; Table 4). Fig. 10 shows a compari-son of the δ13C measured in serpentinites from the Voltri Massif, withthose from the Beigua Unit (western part of the Voltri Massif; Alt et al.,2012), the Atlantis Massif (the basement of the Lost City hydrothermalfield; Delacour et al., 2008) and the Iberian Margin (Schwarzenbach

Table4

Basemen

troc

ksamples:M

ineralog

yan

dcarbon

conten

tsan

disotop

iccompo

sition

sfrom

bulk

rock

analyses.

Samplena

me

Spring

Rock

type

Mineralog

yTo

talc

arbo

n(p

pm)

Totalino

rgan

iccarbon

(ppm

)To

taln

on-carbo

nate

carbon

(ppm

)δ1

3C

TC(‰

)δ1

3 CTIC(‰

)δ1

8 OTIC(‰

)δ1

3C

NCC(‰

)%T

ICof

TC

Type

1:Serpen

tinite

san

dlherzolites

S70_

1S7

0Se

rpen

tinite

serp

±ol

±cp

x+

mgt

+spi+

sulfide

s16

497

67−

13.8

bl.o

.d.

bl.o

.d.

−16

.659

S70_

2S7

0Se

rpen

tinite

serp

±ol

±cp

x+

mgt

+spi+

sulfide

s10

963

46n.a.

bl.o

.d.

bl.o

.d.

n.d.

58S7

0_3

S70

Serpen

tinite

serp

±ol

±cp

x+

mgt

+spi+

sulfide

s17

614

531

−13

.8bl.o

.d.

bl.o

.d.

−23

.582

GOR3

4_3

GOR3

4Lh

erzo

lite

ol+

cpx+

opx+

serp

±am

ph±

chl+

mgt

+spi+

sulfide

s89

654

635

1−

12.9

bl.o

.d.

bl.o

.d.

−18

.461

LER2

0_3_

BLE

R20

Serpen

tinite

serp

+cp

x+

ol±

brc±

clay

minerals+

sulfide

s13

8069

069

0−

7.5

bl.o

.d.

bl.o

.d.

−23

.450

LER2

1_4_

BLE

R21

Serpen

tinite

serp

+cp

x+

ol+

clay

minerals±

Fe-brc

+su

lfide

s38

018

819

2−

12.8

bl.o

.d.

bl.o

.d.

−25

.149

PIO14

_1PIO14

Serpen

tinite

serp

±ol

±cp

x+

mgt

+Cr-spi

+su

lfide

s62

828

434

4−

16.6

bl.o

.d.

bl.o

.d.

−27

.445

Type

2:Serpen

tinite

sBR

1_D2_

A‡

BR1

Serpen

tinite

serp

+cc

+clay

minerals+

amph

+mgt

+su

lfide

s95

2094

5565

−0.2

0.7

−9.5

−27

.199

BR1_

D2_

C1‡

BR1

Serpen

tinite

serp

+cc

+clay

minerals+

amph

+mgt

+su

lfide

s60

0056

8431

6−

0.2

0.3

−9.7

−26

.595

BR1_

D2_

D1‡

BR1

Serpen

tinite

serp

+cc

+clay

minerals+

amph

+mgt

+su

lfide

s18

,390

18,114

276

0.8

1.1

−10

.2−

27.2

98BR

1_D2_

D3‡

BR1

Serpen

tinite

serp

+cc

+clay

minerals+

amph

+mgt

+su

lfide

s86

0586

59bl.o

.d.

0.8

1.3

−8.5

−26

.810

0BR

1_3‡

BR1

Serpen

tinite

serp

+cc

+clay

minerals+

amph

+mgt

+Cr-spi

±ol

±cp

x+

sulfide

s79

4069

5399

0−

1.5

−0.2

−14

.8n.d.

88

Type

2:Metab

asalts

LER2

1_3_

B‡

LER2

1Metab

asalt

alb+

k-fsp+

amph

+ch

l+ep

+ti+

cc74

3064

1410

16−

1.8

−1.0

−24

.2−

6.0

86V18

_3‡

V18

Metab

asalt

bt+

plg+

chl+

ep+

cc+

tlc±

grap

hite

20,615

5641

14,975

−5.6

4.2

−20

.1−

10.3

27

bl.o

.d=

below

limitof

detection;

n.d.

=no

tde

term

ined

;‡=

extens

iveinteractionwithalka

linewater.

237E.M. Schwarzenbach et al. / Lithos 177 (2013) 226–244

et al., 2013). The latter two areas preserve geochemical signatures typicalfor oceanic serpentinites and show a range of δ13CTC = −23.0 to+2.3‰and −25.3 to +1.3‰, respectively. This range is attributed to the pres-ence of variable proportions of inorganic carbon and non-carbonate car-bon (Schwarzenbach et al., 2013). Intense fluid circulation duringmantle exposure on the ocean floor results in the incorporation of largeamounts of seawater carbonate having a δ13C value of ~0 ± 2‰ (e.g.Anderson and Arthur, 1983). Non-carbonate carbon in marine systemsis mainly comprised of organic carbon (TOC) and may representdissolved organic carbon (DOC) incorporated into the serpentiniteduring seawater circulation through the oceanic crust, contributing toδ13CTOC values of −28.9 to −21.5‰ (Delacour et al., 2008). Abiogenicshort-chain hydrocarbons formed by Fischer–Tropsch-Type reactionscould also contribute to the non-carbonate carbon if they subsequentlysorb to the rock (Delacour et al., 2008; Proskurowski et al., 2008). How-ever, due to the volatility of these compounds, it is unlikely that theywould be preserved in the basement rocks.

The close resemblance of the carbon geochemistry of the basementserpentinites to those from oceanic serpentinites suggests that marinecarbonate and marine organic carbon are preserved in the serpentinitesfrom the Voltri Massif. This also implies that inorganic and organic car-bon survives subduction and subsequent obduction onto the continent.Our results agree with findings from Alt et al. (2012) who argue thatsubduction metamorphism of these rocks does not cause significantchanges in the sulfur and carbon geochemistry in the serpentinites.Thus, we conclude that the carbon geochemistry of the Type 1serpentinites represents the original signature from serpentinization onthe ocean floor.

The concentrations of total carbon are at least ten times higher inType 2 basement rocks compared to Type 1 samples (Fig. 5). In Type 2rocks, total carbon ranges between 600 ppm and 2.06 wt.%, with inor-ganic carbon generally contributing between 88 and 100% of the totalcarbon, compared to an average 58% in Type 1 serpentinites (Table 4).This dominance of inorganic carbon in Type 2 rocks is reflected inδ13CTC values of−1.8 to+0.8‰. One exception is themetabasic samplefrom V18 where non-carbonate carbon contributes 73% of the totalcarbon and has a more negative δ13CTC of −5.6‰. Our results and themineralogical observations imply that extensive circulation of the alka-line fluids close to the surface causes precipitation of calcium carbonatein fractures probably facilitated by the transport of atmospheric CO2

throughfissures into the rock. This results in an increase in the inorganiccarbon content of the basement rocks close to the surface. In themetabasic rocks, higher porosities and thus higher fluid fluxes throughthe basement probably cause the overall higher inorganic carboncontents.

The two metabasic samples (V18_3, LER21_3_B), both of whichwere exposed to spring fluids, include some of the highest NCC contentsof 1.5 and 0.1 wt.%, and δ13CNCC of −10.3 and −6.0‰, respectively.Petrological observations indicate that graphite is present in sampleV18_3. Although studies on graphite from various locations reveal arelatively large range in the δ13C with −30.6 to −5.8‰ (Luque et al.,1998), we attribute the less depleted δ13CNCC values and the high NCCcontents to the presence of graphite in these rocks.

In summary, TC, TIC and NCC contents and isotopic compositionsof the serpentinites that can be sampled at the surface demonstratethat 1) the carbon geochemistry of Type 1 serpentinites likelyrepresents the original signatures from serpentinization on theocean floor, suggesting that the carbon within the serpentinites ispreserved and that subduction and subsequent uplift did not impactthe carbon signatures; and 2) during exposure to the alkalinesprings, the inorganic carbon content increases by one order ofmagnitude, which is documented by calcite veins that cut all olderstructures. The variation in TIC contents between Type 1 and Type2 serpentinites also agrees with the alteration mineralogy; i.e. highTIC contents in Type 2 serpentinites coincide with formation of clayminerals and a calcite vein network.

0.0 0.5 1.0 1.5 2.0 2.5 3.0

(Mg+

Fe)

, cat

ions

/7 o

xyge

nsantigorite/lizardite

clinochlore

saponite

montmorillonite

sepiolite

brucite/goethite

0.0

1.0

2.0

3.0

4.0

5.0

6.0

7.0

Si cations / 7 oxygens0.0 0.1 0.2 0.3 0.4 0.5 0.6 0.7 0.8

Al cations / 7 oxygens

0.0

1.0

2.0

3.0

4.0

5.0

6.0

7.0

antigorite/lizardite

clinochlore

saponite

montmorillonite

sepiolite

brucite/goethite

Fig. 4. Selectedmineral compositions (calculated in cations/7 oxygens) of 2 samples:White triangles are froma Type1 serpentinite thatwas collected 2 m from the spring and experiencedminor alteration by alkaline water forming rare thin, fine-grained veins consisting of brucite ± saponite to sepiolite. Gray circles are from a Type 2 serpentinite, where alteration by al-kaline water led to formation of chlorite and fine-grained intergrowths of saponite + serpentine and sepiolite + serpentine. Black stars represent average mineral compositions;black lines are mixing lines between different mineral phases.

238 E.M. Schwarzenbach et al. / Lithos 177 (2013) 226–244

5.3. Evolution of DIC and Ca in the spring waters

Inorganic carbon carried with the spring fluids is important forchemolithoautotrophic communities in the subsurface as the presenceor absence of CO2 in the fluids will directly impact their ability tosurvive. Additionally, the source and availability of CO2 in the subsurfacewill control the amount of carbon available for abiogenic production ofmethane.

The concentrations and isotopic compositions of dissolved inorganiccarbon (DIC) within the spring fluids can be influenced by various pro-cesses: 1) input of atmospheric and soil-derived CO2 to downwellingmeteoric water; 2) removal through water–rock interaction and car-bonate deposition along the flow path; 3) microbial processes in thebasement; 4) input of mantle CO2; 5) abiogenic formation of methane;and 6) late-stage uptake of atmospheric CO2 in the shallow subsurfaceor exit sites before sampling and subsequent to water–rock interac-tions and microbial processes in the basement. An example forextensive late-stage interaction with atmospheric CO2 is provided

S70

_1

S70

_2

S70

_3

GO

R34

_3

LER

20_3

_B

LER

21_4

_B

PIO

14_1

0

5,000

10,000

15,000

20,000 Total inorganic carbon

Total non-carbonate carbon

Type 1 samples:No interaction with alkaline water

Car

bon

cont

ent (

ppm

)

Fig. 5. Total inorganic carbon and total non-carbonate carbon contents of the basement rocks. Tthe serpentinites higher TIC contents also coincide with extensive mineral alteration and form

by spring C11. There the water passes through a 25-meter long, metalpipe (diameter = 3.5 cm) and is exposed to the atmosphere alongthe exit pathway. This results in an average δ13CDIC of −8.7 ± 0.3‰, apH of 10.0, lower Ca and higher Mg concentrations. The high DIC con-tents at spring PIO 14 also suggest that a similar process may be occur-ring in the shallow subsurface. Accordingly, in the following we willfocus mainly on the deeper subsurface processes (processes 2 to 5)and how they may have affected the original DIC composition of thesampled springs.

All of the spring waters are of meteoric origin (Bruni et al., 2002),which is confirmed by the δ18O values of the sampled springs reportedin Table 3. The δ13CDIC of the downwellingfluids is primarily determinedby the δ13C of the atmospheric CO2 at the time the water was lastexposed to the atmosphere and the contribution of CO2 from organicmatter respiration in the soils. The average δ13CDIC of groundwatersand rivers can vary from −5 to −22‰ (e.g. Waldron et al., 2007). TheDIC of two rivers measured in the Voltri Massif lie within this range,with δ13CDIC values of −15.6 and −11.2‰ and DIC concentrations of

BR

1_D

2_A

BR

1_D

2_C

1

BR

1_D

2_D

1

BR

1_D

2_D

3

BR

1_3

LER

21_3

_B

V18

_3

Type 2 samples:Extensive interaction with alkaline water

ype 2 samples have significantly higher inorganic carbon contents than Type 1 samples. Ination of clay minerals.

10 100 1,000

DIC (μmol/L)

-25.0

-20.0

-15.0

-10.0

-5.0

0.0

5.0

δ13C

DIC

Evolution from riv

er water

to Ca-OH waters

BR watersBR 2 river GOR waters

L43 waters

L43 river

LER watersPIO waters

Fig. 6. Concentration (log scale) and isotopic composition of the DIC of the sampledsprings in the Voltri Massif. The two rivers (black symbols) have significantly higher DICcontents. Concurrent closed-system serpentinite dissolution and carbonate precipitationin the basement leads to a depletion in DIC and low δ13CDIC values and promotes increasesin Ca and the acquisition of Ca–OH compositions. The trend to higher δ13CDIC valuesassociated with a slight decrease in DIC (black arrow) could be the result of DIC removalby microbial activity in the subsurface or abiotic production of methane.

-30.0 -25.0 -20.0 -15.0 -10.0 -5.0

-25.0

-20.0

-15.0

-10.0

-5.0

0.0

5.0Ligurian springsOmanWestern USA

δ13Ccarb (‰)

δ18O

carb (

‰)

Fig. 7. Isotopic compositions of the carbonates sampled at the alkaline springs in the VoltriMassif plotted as δ18O against δ13C of the carbonate. Similar isotopic compositions character-ize carbonates that are associated with alkaline springs in the Oman (Samail ophiolite; datafrom Clark et al., 1992), and in the Western USA (data from O'Neil and Barnes, 1971).

239E.M. Schwarzenbach et al. / Lithos 177 (2013) 226–244

2122 and 930 μmol/L, respectively (Table 3; Fig. 6). Successive interac-tion of the meteoric water with the basement rock and the evolutiontowards closed system conditions modifies the δ13CDIC value of thefluids and removes carbon through carbonate deposition (Bruni et al.,2002). Simple precipitation of calcium carbonate in a closed systemcan be modeled by a Rayleigh fractionation process and will lead to de-creasing Ca-concentrations and depletion in 13C in the fluid because thefractionation between calcite and dissolved bicarbonate is ~2‰ (Deineset al., 1974). Assuming an initial δ13CDIC of −15‰, a concentration of2 mmol (corresponding to the BR2 river) and a fractionation factor of2‰, 90% removal of DIC through carbonate precipitation will result ina final δ13CDIC of approximately −24‰. The DIC concentrations of thesprings investigated here range from 8.8 to 15.6 μmol/L and thus reflect90–95% removal of DIC relative to the concentrations in the two rivers.However, these springs show characteristically high δ13CDIC values thatcannot solely be attributed to closed-system carbonate precipitation,and 2–3 times higher Ca concentrations than the two rivers (Table 3;see also Cipolli et al., 2004). The Ca-concentrations are consistentwith reaction path modeling with Ca being continuously suppliedto the system by dissolution of serpentinite under closed-systemconditions with respect to CO2 and concomitant carbonate precip-itation (Bruni et al., 2002). Only the spring waters from BR3 andGOR34 have δ13CDIC values that can be attributed to Rayleigh frac-tionation by carbonate precipitation, which suggests that otherprocesses have influenced the carbon isotope signatures and DICconcentrations at other springs.

Close examination of the concentration data and C-isotope composi-tions of the alkaline springs reveals a linear relationship between DICconcentrations and δ13CDIC values (Fig. 6), whereby large variations inCa concentrations of ~400 μmol/L are associated with minimal varia-tions in DIC concentrations (Table 3). An increase in δ13CDIC of nearly20‰ is associated with a decrease in DIC of ~5 μmol/L (Fig. 6) andsuggests that this decrease is related to a process that imparts a largeisotopic fractionation of ~45‰. Such large fractionations could reflectmicrobial activity in the basement and consumption of DIC throughmethanogenesis, where continuous removal of 13C-depleted DIC fromthe fluid results in an enrichment of 13C of the remaining DIC. At suchlow DIC concentrations, even very small degrees of methanogenesiscould affect the isotopic compositionswithout producingnotable effectsin the methane concentrations of the fluids.

Alternatively, abiotic reduction of CO2 and production of methanethrough Fischer–Tropsch type reactions could also influence the compo-sitions in a similar way. Thermodynamic equilibrium between CO2 andCH4 predicts that methane is strongly depleted in 13C, resulting in highδ13C values of the residual CO2 in the fluid that could contribute to thetotal DIC of the system. In experiments, δ13CCH4 values as low as−50‰ and δ13CCO2 values between −0.8 and +20.8‰ have beenfound (Horita and Berndt, 1999; McCollom and Seewald, 2006).

Similarly, leaching of fluid inclusions in the basement rock couldpotentially release mantle CO2 into the fluids. Such an effect has beenfound as a result of extensive hydrothermal alteration of oceaniccrust and can additionally influence the DIC composition (Kelley andFrüh-Green, 1999). Mantle carbon has an average δ13C of ~−5‰(e.g. Deines, 2002; Ohmoto and Goldhaber, 1997; Schulze et al., 1997).Leaching of fluid inclusions in the basement rocks could thereforelead to higher δ13CDIC values, butwould result in a concomitant additionof C to the system. This process could only be envisioned at GOR34,where lherzolites form the basement of the spring waters. The othersprings are hosted by serpentinites, which have already been extensivelyaltered during serpentinization on the ocean floor, and are unlikely toretain fluid inclusions.

In summary, primarily two processes are consistentwith the presentdata set: 1) Continuous precipitation of calcite leads to C-uptake and lowδ13CDIC values of the residual DIC, and is associatedwith an increase in Cadue to closed-system serpentinite dissolution, 2) removal of DIC bymicrobial activity (e.g., methanogenesis) resulting in a 13C-enrichmentin the residual fluid. Although abiotic methane formation could producesimilar isotopic signatures of the final DIC and catalysts such as awaruiteand heazlewoodite are present in the basement rocks (Schwarzenbach,2011), it may be less likely due to the low temperature conditions inthese systems. Similarly an input of mantle CO2 is less likely in theserpentinites, while it cannot entirely be excluded for the springs thatare situated within the lherzolites. We thus interpret the high-pH,Ca-rich springs to represent highly evolved waters that have nearly losttheir capacity to precipitate carbonate in the subsurface and in whichmicrobial activity may be important in controlling the final C-isotopecomposition of the residual fluids. In addition, the resulting high pH ofthese springs result in a further capacity to sequester atmospheric CO2

and precipitate carbonate at the exit sites.

LER20_A

LER20_B

LER20_C

spring

LER20_2

LER20_1

-30.0 -25.0 -20.0 -15.0-20.0

-15.0

-10.0

-5.0LER20_1

LER20_2

LER20_A/B/C

δ13CTIC (‰)

δ18O

TIC

(‰

)

increasing distance from the spring

BA

Fig. 8. Variation of the isotopic composition of the carbonate deposits sampled at spring LER20, where five carbonate samples were takenwith distance from the spring outlet. A) SampleLER20_1 taken at the edge of the pool that formed underneath the pipe; Sample LER20_2 is soaked by alkaline water that exits along fractures above the pipe; Samples LER20_C, LER20_B,and LER20_Awere takenwith increasing distance from the spring (1.2 m, 2.5 m, and ~4 m from the spring, respectively). B) Sample LER20_1 (gray triangles) from underneath the metalpipe shows the strongest depletions in 13C and 18Owith δ13C as low as−26.1 and δ18O as low as−17.7‰. With increasing distance from the spring, the depletion in 13C and 18O decreases(gray arrow) resulting in values of −17.9 and −5.9‰ for δ13C and δ18O, respectively, in sample LER20_B. Multiple data points for samples LER20_1 and LER20_2 indicate differentextracted layers of the carbonate deposit.

240 E.M. Schwarzenbach et al. / Lithos 177 (2013) 226–244

5.4. Formation of the carbonate deposits

The extremely low DIC contents of the springs suggest that thecarbonate deposits incorporate atmospheric CO2 during precipitation.Carbonates with strongly negative carbon and oxygen isotope composi-tions are ubiquitously found where highly alkaline spring watersemerge from peridotites or serpentinites (Clark et al., 1992; O'Neiland Barnes, 1971). Clark et al. (1992) found δ13CTIC values of as low as−27.5‰ and δ18OTIC values of −16.9 to −14.0‰ (PDB) in carbonatedeposits in the Oman ophiolite (Fig. 7) and attributed them to extremekinetic fractionation during CO2 hydroxylation. To explain their datathey proposed a thin-film model (after Liss, 1973; Quinn and Otto,1971; Usdowski and Hoefs, 1986), which is illustrated in Fig. 11 andinvolves the following sequence of reactions:

CO2ðgÞ⇔CO2ðaqÞ ð4Þ

-28

-24

-20

-16

-0.5

0.0

0.5

1.0

1.5

2.0

2.5

3.0

3.5

4.0

4.5

Dis

tanc

e fr

om s

prin

g (in

m)

-20

-15

-10 -5 6 8

Na δ18OTIC (‰)δ13CTIC (‰)

Fig. 9. Carbon and oxygen isotope composition of the inorganic carbon of the carbonate depwhich was sampled along a profile with increasing distance from the spring shown in Fig. 8A.circles = samples taken with distance from the spring.

CO2ðaqÞ þ OH−⇔HCO

−3 ð5Þ

Ca2þ þ CO

2−3 ⇔CaCO3 ð6Þ

Clark et al. (1992) demonstrated that CO2 carboxylation (Eq. (5)) ismuch slower than diffusion of atmospheric CO2 into thewater (Eq. (4)),and thus is the process controlling isotope fractionation mainly byaqueous kinetic effects. In experiments, Clark et al. (1992) simulatedthe conditions at which the travertines in the Oman ophiolite form. Byprecipitating barium carbonate (BaCO3) they determined a carbonisotope fractionation factor of Δ13CCO2(g)–BaCO3 = −15.5‰, which canalso be used for calcium carbonate.

The carbonate samples from the Voltri Massif give a similar fraction-ation between atmospheric CO2 and carbonate. Compared to a δ13Cvalue of −8.5‰ for atmospheric CO2 (Longinelli and Selmo, 2006) thesampled carbonates yield an average depletion of −11.9‰. However,samples characterized by more or less instant precipitation (indicated

10 12 14 16

(mg/l)

0 5 10 15 20 25

Mg (mg/l)

20 25 30 35 40

Ca (mg/l)

osits sampled at spring LER20, and concentration of Na+, Mg2+, and Ca2+ in the water,Triangles = spring water and carbonates, sampled directly underneath the spring outlet;

Table 5Water chemistry and isotopic composition of the carbonates sampled at spring BR1.

Rock sample δ13C TIC (‰) δ18O TIC (‰) Water sample pH T (°C) Na (mg/L) Mg (mg/L) Ca (mg/L) δ13C DIC (‰) δ18O H2O (‰) VSMOW

BR1_1_L2 −27.2 −17.8 BR1 11.96 14.8 23.48 bl.o.d. 44.25 1.3 −6.8BR1_1_L3 −26.7 −17.6BR1_5a −21.7 −14.1 BR1_5 (pool) 11.68 14.9 18.55 2.50 15.62 −6.7

a Sample BR1_5 was taken approx. 5 m from the spring outflow of spring BR1.

241E.M. Schwarzenbach et al. / Lithos 177 (2013) 226–244

by the subscript “c” in Table 2) have a depletion of−15.9‰ and thus arewithin the range of the experiments of Clark et al. (1992).

The δ18O of the carbonates in the Voltri Massif show a largervariation between −18.7 and −2.8‰. Similar to carbon, fractionationof the oxygen isotopes is dominated by kinetic rather than equilibriumisotopic effects. Oxygen isotope values depend on temperature andthe δ18O composition of both the water and the atmospheric CO2, andcan be calculated as follows (after Clark et al., 1992):

δ18OBaCO3 ¼ 1=3ðδ18OH2O þ ε18OOH––H2OÞ þ 2=3ðδ18OCO2ðgÞ � ε18OCO2ðgÞ–CO2ðaqÞÞð7Þ

According to Eq. (7), assuming a δ18OCO2 = +11.0‰, ε18OOH– –H2O =−40‰, and ε18OCO2(g) – CO2(aq) = +2‰, at 25 °C (values from Clarket al., 1992) and using δ18OH2O values measured for the individual springwaters (Table 3), the carbonates from theVoltriMassif should have an av-erage δ18OCaCO3 value of approximately−19.7‰. However, themeasuredδ18OCaCO3 values are slightly less depleted with an average of −13.4‰for all analyzed samples, and −17.2‰ for the carbonates that havebeen precipitated immediately underneath the spring outflow.

5.4.1. Fluid evolution with distance from the spring and influence of the pHThe carbon and oxygen isotope compositions show strong

correlations with distance from the spring (Fig. 8) and therefore,whether precipitation of the carbonate occurred immediately orafter some time of equilibration with the atmosphere. At springBR1, two samples with variable distance to the spring were takento compare the fractionation factor: one directly underneath the

0 2 4 6 8 10

TC (%C)

-30

-25

-20

-15

-10

-5

0

5

0 2 4 6 8 10

TIC (% C)

0

5,00

0

10,0

00

TOC (ppm)

δ13C

TC (

‰)

Fig. 10. Carbon contents and isotope signatures of basement rocks of the Voltri Massif (Type 1BeiguaUnit (data fromAlt et al., 2012) and oceanic serpentinites from theAtlantisMassif along(data from Schwarzenbach et al., 2013).

pipe (BR1_1) and a second one 5 m from the spring (BR1_5), takenat a small barrier between two pools (Table 5). There the pH is 11.7and despite its distance from the spring, both carbon and oxygen iso-topes are strongly negative (δ13CTIC = −21.7, δ18OTIC = −14.1‰;Table 5). In contrast, at spring LER20 δ13C and δ18O are less negative(δ13CTIC = −18.6‰, δ18OTIC = −6.7‰) at 1.2 m from the spring,where the pH has decreased to 9.4. As illustrated in Fig. 11, theisotopic fractionation of the carbonates depends on the dominant re-action during carbonate precipitation (Eqs. (4), (5), and (6)), whichis controlled by the pH (Clark et al., 1992). At high pH (pH N7)hydroxylation of CO2 (Eq. (5)) dominates while at low pH (pH b7)hydration of CO2 dominates and induces a distinctly lower fraction-ation factor (Clark et al., 1992). Importantly, interaction with theatmosphere and invasion of CO2 leads to a decrease in pH (Eq. (8)).

CO2ðaqÞ þ H2O ¼ HCO−3 þ H

þ ð8Þ

The flow rate, the residence time in a pool or the surface of interac-tion with the atmosphere (including turbulent mixing during waterflow) are all factors that control CO2 exchange between the atmosphereand the fluids, and thus control the change in pH of the water. Conse-quently, the extent of equilibration with atmospheric CO2 ultimatelyinfluences the overall fractionation during carbonate precipitation.Additional factors are the mixing with rainwater, which influences theisotopic composition of CO2 (aq) and H2O, and the pH. Seasonal varia-tions are likely the cause for the isotopic differences within differentlayers of individual samples.

15,0

00 -4 -2 0 2 4 6-2

5-2

0-1

5-1

0 -5 0 5-3

0

-25

-20

-15

-10 -5

δ13CTIC (‰) δ13CNCC (‰) δ18OTIC (‰)

Voltri Massif (Type 1 samples)Voltri Massif (Type 2 samples)Beigua UnitIberian Margin serpentinitesAtlantis Massif serpentinites

and Type 2 samples), compared to carbon signatures measured in serpentinites from thetheMid-Atlantic Ridge (data fromDelacour et al., 2008) and the continental IberianMargin

CO2 (g) (δ13CCO2 ≈ -8.5‰)

CO2 (aq) CaCO3HCO3- CO3

2-

OH-

OH-

OH-

Ca2+

Ca2+H2O

atmosphere

solution thin film

solution thin film

bulk solution

Eq. 4

Eq. 5 Eq. 6

Fig. 11. Sketch of the thin-film model after Clark et al. (1992), showing the reactionsequence during carbonate precipitation from an alkaline fluid. The reactions arenumbered according to the equations in the text. Hydroxylation of CO2 is the processthat controls isotopic fractionation.

242 E.M. Schwarzenbach et al. / Lithos 177 (2013) 226–244

5.5. Carbon sequestration during fluid evolution

This study shows that interaction ofmeteoricwaterwith serpentinitesand peridotites results in the storage of carbon, most evident by theformation of carbonate deposits at the surface and carbonate withinfractures of the sampled basement rocks. While the carbonate withinthe surface of the basement could significantly add to the total car-bon sequestered within this system (up to 2 wt.%), field observationsof the springs indicate that the amount of CO2 sequestered long-termthrough surficial carbonate precipitation is probably small, as a re-sult of periodic dissolution of the deposits. However, the mainprocess that results in long-term carbon sequestration is the removalof carbon from the meteoric waters during their evolution fromMg–HCO3 to Ca–OH waters – the latter of which are strongly depletedin inorganic carbon – and its storage in the basement rocks.

Bruni et al. (2002) measured average maximum DIC concentrationsof the Mg–HCO3 waters of around 3.27–4.92 mmol/L (200–300 mg/Ldominated by HCO3

− see Fig. 12; individual data reported in Marini

Average maximum DICconcentration

Mg-HCO3waters Ca-OH

waters

Closed system

evolution

and carbon removal

5.0 6.0 7.0 8.0 9.0 10.0 11.0 12.0pH

1

10

100

1,000

10,000

DIC

(μm

ol/L

)

Fig. 12. Evolution of the Mg–HCO3 waters to Ca–OH waters showing the spring waters(gray circles) and the two sampled rivers (black circles) in comparison to the Mg–HCO3

and Ca–OH waters analyzed by Bruni et al. (2002) (light gray fields). The dark gray barrepresents estimated average maximum DIC concentrations of around 200–300 mg/Lreported by Bruni et al. (2002) and Marini and Ottonello (2002). The dark gray arrowillustrates the evolution from open to closed system conditions with the formation ofCa–OH waters and removal of almost the entire carbon pool in the fluid.

and Ottonello, 2002). As discussed above, the compositions of theCa–OH waters indicate that almost the entire carbon is removed fromthe fluid; we measured as little as 8.8 μmol/L (Table 3). Consequently,this suggests that almost the entire inorganic carbon pool present inthe original fluids, i.e. up to 4.9 mmol/L, is likely sequestered along theflow path. To model the extent of carbon sequestration through carbonremoval duringwater–rock interaction in the Voltri Massif, we assumedan annualwater input of 1000–1100 L/m2 per year (= 1000–1100 mmas rainfall; annual rainfall for Genoa is ~1073 mm, we assume that theaverage rainfall is constant within that range) and a drainage basinarea of 760 km2 (from Ferraris et al., 2012, representing the approxi-mate basin area of the rivers draining the Voltri Massif). The amountof rainwater that reacts with the serpentinite and peridotite basementrocks of the Voltri Massif depends on the run off rate. This rate deter-mines what percentage enters the groundwater system and eventuallythe basement rocks. Fig. 13 shows the calculations for the determinedvalues of catchment area and rainfall as a function of runoff. Consideringa range of total carbon available for carbonate precipitation between3.27 and 4.92 mmol/L and a runoff rate between 50 and 80% (i.e. 20 to50% of the rain water will enter the basement), we calculate a potentialfor carbon sequestration of 0.50 to 2.05 × 109 mol/year. Similarly,Kelemen and Matter (2008) calculated that through weathering of theOman ophiolite ~104 to 105 tons CO2/year (b2.27*1010 mol/year) areconverted to carbonate. For such estimates, the rainfall amount in thearea of exposed mantle rock, and thus how much water reacts withthe mantle rocks, play an essential role. Low rainfall at the Samailophiolite likely decreases the amount of carbon that is sequesteredthan if water–rock ratios would be higher. The Voltri Massif is ~20times smaller than the exposed ophiolite sequence in the Oman(Kelemen and Matter, 2008). This suggests that the Voltri Massif has ahigher annual carbon sequestration rate per square meter than theophiolite in the Oman.

Cipolli et al. (2004) have modeled the sequestration of CO2 throughhigh-pressure injection into the deep aquifers and emphasized theimportance of maintaining sufficient porosity. Mineral precipitationof e.g. magnesite, serpentine, or amorphous silica during water–rockinteraction significantly lowers the porosity of the aquifers and mayslacken or even hinder effective CO2 sequestration. Thus, artificialhydrofracturing and amplification of the porosity may be the main pro-cess to efficiently sequester CO2 in variably serpentinized peridotites(Kelemen and Matter, 2008). Additionally, it needs to be consideredthat the global anthropogenic CO2 emissions are about 30 billion tons

2.05*109 mol/yr

4.97*108

mol/yr

0 10 20 30 40 50 60 70 80 90 100

Run off rate

0

1*109

2*109

3*109

4*109

5*109

Tot

al u

ptak

e of

car

bon

(in m

ol/y

r) Average rainfall: 1100 mm/yr

Average rainfall: 1000 mm/yr

Fig. 13. Calculation modeling the potential of the Voltri Massif basement rocks to seques-ter carbon during fluid evolution to the Ca–OHwaters. We assume an annual water input(as amount of rainwater) of 1.0–1.1x103 L*m−2*yr−1 and a catchment area of 760 km2.Assuming a run off rate of 50–80% we calculate that annually 0.49–2.05 × 109 mol C isincorporated into the basement rocks.

243E.M. Schwarzenbach et al. / Lithos 177 (2013) 226–244

of CO2 every year. Thus, natural CO2 sequestration in ultramafic com-plexesmust be considered as of minor importance in reducing the glob-al atmospheric CO2.

6. Summary and conclusions

This study has shown that serpentinites recycled through subduc-tion and subsequent uplift onto continents play an important role inthe cycling of carbon, affecting various biological and abiologicalprocesses. Interaction of the serpentinites with meteoric water leadsto the formation of alkaline fluids that have the capacity to sequesterCO2 as a consequence of closed system fluid–rock interaction andthrough carbonate precipitation on the surface orwithin the uppermostzones of the serpentinites. Interaction with alkaline fluids results inthe formation of fine-grained saponite and sepiolite and possibly ser-pentine. The mineralogy of the serpentinites and the abundance ofcarbonate deposits along fractures of the basement show that cracksand fractures serve as important fluid-pathways for meteoric water tocirculate and produce the Ca-rich, alkaline fluids.

The overall large range in δ13CDIC values implies that the DIC issensitive to subsurface and surface processes and that the final DICsignatures reflect a combination of processes. The concentration andisotopic composition of the DIC in the spring waters provides evidencefor 1) the precipitation of calcium carbonate following a Rayleighfractionation model and 2) removal of DIC by microbial activity in thesubsurface.

Kinetic fractionation effects dominate the isotopic composition ofthe carbonate deposits, resulting in strongly negative δ13C and δ18Ovalues as is typical for precipitation fromhighly alkaline fluids. Samplingof carbonates from various settings showed that the isotopic composi-tion is controlled by the extent of equilibration with atmospheric CO2,which in turn is controlled by fluid flow rate, temperature, and time ofequilibrationwith air.With increasing distance from the spring outflow,fractionation decreases due to a decrease in pH, eventually leading toequilibrium isotopic compositions. Periodic dissolution of carbonate de-posits, however, suggests that the importance of carbonate formationon the surface may have a lesser importance in the long-term carbonsequestration. In contrast, comparison of spring-water altered andunaltered basement rocks demonstrates that the evolution towardsalkaline spring water results in a tenfold increase in carbon contentwithin the serpentinites. The formation of alkaline fluids results in theincorporation of up to 2 wt.% C, which suggests that the basementrocks of the Voltri Massif have a significant capacity to incorporate car-bon during fluid–rock interaction. We present model calculations thatsupport the conclusion that the basement rocks in the Voltri Massifplay an essential role in the active, long-term sequestration of carbon,but that compared to the global anthropogenic CO2 emissions the naturalsequestration of CO2 in the Voltri Massif is of minor importance.

Acknowledgments

Wewould like to thankMaria Coray-Strasser, Stewart Bishop, FannyLeuenberger and Chantal Ulmer for their help with analyses and MarcoMolinari and Luigi Marini from the University of Genoa for their help inthe field. We also would like to thank J. Alt and C. Boschi for the helpfulcomments that greatly improved the manuscript and M. Scambelluri isthanked for editorial handling. This work was supported by Swiss NSFgrants no. 200020-116226 and no. 200020-124669 to Früh-Green.

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