Hydrothermal processes in partially serpentinized ... · lower section of an ancient hydrothermal system, where conditions were highly reducing and water–rock ratios very low. Thus,

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Contrib Mineral Petrol (2014) 168:1079DOI 10.1007/s00410-014-1079-2

ORIGINAL PAPER

Hydrothermal processes in partially serpentinized peridotites from Costa Rica: evidence from native copper and complex sulfide assemblages

Esther M. Schwarzenbach · Esteban Gazel · Mark J. Caddick

Received: 5 March 2014 / Accepted: 23 October 2014 © Springer-Verlag Berlin Heidelberg 2014

formed by local addition of a hydrothermal fluid that likely interacted with adjacent mafic sequences. We suggest that the peridotites today exposed on Santa Elena preserve the lower section of an ancient hydrothermal system, where conditions were highly reducing and water–rock ratios very low. Thus, the preserved mineral textures and assemblages give a unique insight into hydrothermal processes occur-ring at depth in peridotite-hosted hydrothermal systems.

Keywords Native copper · Sulfides · Peridotite · Serpentinization · Santa Elena Ophiolite

Introduction

Serpentinization is a widespread process that is found where ultramafic rocks react with seawater, hydrothermal fluids or metamorphic fluids within subduction zones (e.g., Cannat et al. 1992; Hyndman and Peacock 2003; Mével 2003; Früh-Green et al. 2004; Cannat et al. 2010). Dur-ing reaction of water with the primary minerals olivine and pyroxene, H2 is formed due to oxidation of Fe2+ to Fe3+ (e.g., Frost 1985; Bach et al. 2006). As a result, highly reducing conditions are produced that are rarely seen in other geological environments. These high H2 conditions allow the stability of native metals, Fe–Ni alloys (e.g., awaruite, taenite) and other rare sulfides such as heazle-woodite or polydymite (Frost 1985; Klein and Bach 2009). Despite this seemingly hostile environment, serpentiniza-tion has been shown to provide the necessary energy source for microbial activity and peridotite-hosted hydrothermal systems have been found to host diverse microbial com-munities (Kelley et al. 2005; Brazelton et al. 2006; Russel et al. 2010; Brazelton et al. 2011), making these environ-ments of great interest for studying processes that link the

Abstract Native metals and metal alloys are common in serpentinized ultramafic rocks, generally representing the redox and sulfur conditions during serpentinization. Vari-ably serpentinized peridotites from the Santa Elena Ophi-olite in Costa Rica contain an unusual assemblage of Cu-bearing sulfides and native copper. The opaque mineral assemblage consists of pentlandite, magnetite, awaruite, pyrrhotite, heazlewoodite, violarite, smythite and copper-bearing sulfides (Cu-pentlandite, sugakiite [Cu(Fe,Ni)8S8], samaniite [Cu2(Fe,Ni)7S8], chalcopyrite, chalcocite, bornite and cubanite), native copper and copper–iron–nickel alloys. Using detailed mineralogical examination, electron micro-probe analyses, bulk rock major and trace element geo-chemistry, and thermodynamic calculations, we discuss two models to explain the formation of the Cu-bearing mineral assemblages: (1) they formed through desulfurization of primary sulfides due to highly reducing and sulfur-depleted conditions during serpentinization or (2) they formed through interaction with a Cu-bearing, higher temperature fluid (350–400 °C) postdating serpentinization, similar to processes in active high-temperature peridotite-hosted hydrothermal systems such as Rainbow and Logatchev. As mass balance calculations cannot entirely explain the extent of the native copper by desulfurization of primary sulfides, we propose that the native copper and Cu sulfides

Communicated by O. Müntener.

Electronic supplementary material The online version of this article (doi:10.1007/s00410-014-1079-2) contains supplementary material, which is available to authorized users.

E. M. Schwarzenbach (*) · E. Gazel · M. J. Caddick Department of Geosciences, Virginia Tech, 4044 Derring Hall, Blacksburg, VA 24061, USAe-mail: [email protected]

http://dx.doi.org/10.1007/s00410-014-1079-2

Contrib Mineral Petrol (2014) 168:1079

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geochemical cycles between the lithosphere, hydrosphere and biosphere (Früh-Green et al. 2004; Schwarzenbach et al. 2012, 2013b).

Variably serpentinized peridotites and their sulfide and oxide assemblages have been studied in ultramafic bod-ies tectonically emplaced on continents (Eckstrand 1975; Garuti et al. 1984; Peretti et al. 1992), along mid-ocean ridges, where detachment faulting causes exposure of ultra-mafic rocks to seawater inducing extensive serpentinization (Bach et al. 2004; Alt et al. 2007; Delacour et al. 2008a, b), in fossil peridotite-hosted hydrothermal systems (Hop-kinson et al. 2004; Schwarzenbach et al. 2013b), within the mantle wedge or the subducting plate (Hyndman and Pea-cock 2003; Alt and Shanks 2006; Scambelluri and Tonarini 2012) and in cratonic lithospheric mantle xenoliths (Lorand and Gregoire 2006). Typical primary sulfides in peridotites are pentlandite ± pyrrhotite ± chalcopyrite and occur as inclusions within silicates (e.g., Lorand 1989a, b). Only rarely are native metals found in peridotites, while exten-sive Cu–Fe–Ni sulfides are usually associated with seafloor hydrothermal systems or are exposed on the continent as volcanogenic massive sulfide (VMS) ore deposits. Spe-cifically, Cu-rich sulfide assemblages have been related to hydrothermal leaching of mafic sequences, but native cop-per has also been related to alteration of primary Cu-bear-ing sulfide minerals or even of primary mantle origin (e.g., Abrajano and Pasteris 1989; Tsushima et al. 1999).

The opaque mineralogy in serpentinized peridotites records the hydrogen/oxygen and sulfur fugacities during serpentinization reactions (Frost 1985; Klein and Bach 2009). While initial serpentinization allows the stability of native metals and metal alloys, completely serpentinized peridotites typically preserve high-sulfur assemblages and magnetite or hematite (Eckstrand 1975; Alt and Shanks 1998; Delacour et al. 2008a; Schwarzenbach et al. 2012). Thus, the study of the opaque mineral assemblages is key to understanding the evolution of the hydrogen and sulfur fugacities during the serpentinization process. Addition-ally, serpentinites play an important role in many global geochemical cycles and control the transport of various species (e.g., H2O, sulfur) into the mantle (e.g., Scambel-luri et al. 1995; Ulmer and Trommsdorff 1995; Scambelluri and Tonarini 2012; Alt et al. 2013). Revealing the processes that accompany serpentinization is therefore required to completely characterize the geochemical cycling between Earth’s surface and Earth’s mantle.

On the Santa Elena peninsula in Costa Rica, variably serpentinized peridotites crop out together with layered and pegmatitic gabbros and are intruded by mafic dikes (Gazel et al. 2006). Due to the low degree of serpentinization and low degree of weathering in most samples, different stages of serpentinization and various sulfide textures are well preserved. This makes the Santa Elena peridotites ideal

to study the successive processes that are associated with the hydration of peridotites. Moreover, the discovery of the presence of native copper together with a very diverse sulfide mineralogy in the samples from the Santa Elena Ophiolite is of particular interest in understanding both the redox and the sulfur conditions during serpentinization, and the possible interaction with high-temperature hydrother-mal fluids (>350 °C) pre- or postdating lower temperature serpentinization (~200–250 °C). Here, we present data on the opaque mineralogy, sulfide and metal mineral chemis-try and bulk rock chemistry of the Santa Elena peridotites with the goal of evaluating the source and speciation of the copper-bearing assemblages and to give insights into the hydrothermal evolution of these peridotites.

Geological setting and sample selection

The Santa Elena Ophiolite is located on the west coast of Costa Rica and comprises an area of 250 km2 of mafic and ultramafic lithologies (Fig. 1; Gazel et al. 2006). Geotec-tonically, Costa Rica is today situated on the triple junction of the Cocos, Caribbean and Nazca Plates. Along the Mid-dle American Trench, the Cocos Plate is being subducted underneath the Caribbean Plate, resulting in an active vol-canic front, while along the pacific side of Costa Rica sev-eral oceanic complexes have been accreted onto the Car-ibbean Plate (Hauff et al. 2000; Denyer and Gazel 2009; Herzberg and Gazel 2009 and references therein). The Santa Elena peridotites have generally been correlated with peridotites cropping out along the Costa Rica–Nicaragua border, suggesting an E-W fossil suture zone between dif-ferent tectonic blocks (Tournon et al. 1995). The Santa Elena Ophiolite is locally covered by reef limestones of Campanian age, suggesting that the Santa Elena Peninsula was emplaced during the Upper Cretaceous with a peri-dotitic complex at the hanging-wall and an igneous-sedi-mentary complex at the footwall—the Santa Rosa Accre-tionary complex (Baumgartner and Denyer 2006; Denyer and Gazel 2009). The Santa Elena Nappe (Fig. 1) contains variably serpentinized peridotites, dunites and locally lay-ered gabbros. Various generations of pegmatitic gabbros and diabase dikes cut the peridotites. Some of these dikes do not preserve chilled margins, suggesting that they were emplaced into a hot mantle host preceding serpentiniza-tion (Gazel et al. 2006). A secondary mineralogy in the mafic lithologies composed of albite + epidote + actino-lite + chlorite has been ascribed to ocean floor metasoma-tism (Gazel et al. 2006).

The samples studied here were collected at various locations within the Santa Elena Ophiolite (Fig. 1) and include boulders within streams that were collected due to great preservation conditions. The samples include


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lherzolites, harzburgites and dunites with variable degrees of serpentinization.

Methods

The mineralogy and petrology of the peridotites were ini-tially studied in thin section with transmitted and reflected light microscopy. The mineral chemistry of the sulfides was determined on a Cameca SX-50 electron microprobe (EMP) at 15 kV acceleration potential, 20 nA current and 1 μm beam size, using natural and synthetic mineral stand-ards. Relative analytical error is better than 1 % (1σ) except for element contents <1 wt%, where the analytical error is better than 4 % (1σ). Element distribution maps of selected grains and areas were collected using the EDS (energy dis-persive spectrometer) system and were run between 2 and 12 h and at a current of 40–100 nA, depending on run time.

Bulk rock samples were powdered using an alumina mill and were fluxed into homogeneous glass disks with ultrapure Li2B4O7 from Spex® (certified ≪1 ppm blank for all trace elements) at the Petrology Lab at Virginia Tech for X-ray fluorescence (XRF) and inductively coupled plasma mass spectrometry (ICP-MS) analyses. Major ele-ments were collected following the methods described in Mazza et al. (2014). The analytical error for 10 replicates of BHVO-2g was <1.5 % for all major elements. Trace ele-ments were collected from the same fluxed glasses with an Agilent 7500ce ICPMS coupled with a Geolas laser abla-tion system, with a He flow rate of ~1 L/m−5 Hz and an energy density on sample ~7–10 J/cm2, following the pro-cedures detailed in Mazza et al. (2014). Data were cali-brated against USGS standards BHVO-2g, BCR-2g and BIR-1g, using Si from XRF as an internal standard and the standard element values reported in Kelley et al. (2003). The analytical error for 10 replicates of BHVO-2g was

N

Potrero Grande tectonic window

Punta El Respingue

Layered gabbros (124 Ma)

Playa Santa Rosa

Islas Murciélago

Pacific Ocean

5 km

Santa Elena Nappe

Pillow and massive basalts (109 Ma)

Dolerite dikes Santa Elena thrustFaults

Dike swarm

Santa Rosa Accretionary Complex

10°55’

10°50’

85°50’85°55’85°60’ 85°45’ 85°40’

SE_P9SE_P7

SE_P5

SE_P3

SE10_01, 02

SE_P1, P2

SE_P10

SE10_12 SE10_16

SE10_09SE10_05, 06

SE_P4

SE10_19SE_P8

Costa Rica

Fig. 1 Geological map of the Santa Elena Ophiolite in Costa Rica with the location of the analyzed peridotite samples (after Gazel et al. 2006)


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<5 % for all elements with the exception of Ni, Cu and Yb (<7 %).

Results

Mineralogy

The samples studied here are lherzolites, clinopyroxene-rich harzburgites and dunites with a degree of serpentini-zation ranging between 30 and 100 % (Table 1). Hydra-tion led to replacement of olivine by serpentine forming a typical mesh texture. In all samples, olivine replacement is more advanced than decomposition of pyroxene, where ini-tial replacement by serpentine typically occurs along frac-tures within the grains. The serpentinization textures show a mesh texture to rare ribbon veins and parallel veining, indicating serpentinization under static conditions. Serpen-tinization of olivine resulted in the formation of serpen-tine + brucite veins, with almost pure brucite present in the center of some veins. Hydration also led to the formation of minor amounts of chlorite and amphibole.

The opaque mineral associations include pentlandite, magnetite, awaruite, pyrrhotite, heazlewoodite, violarite and smythite (Table 1; Fig. 2). Numerous peridotite sam-ples contain copper sulfides (chalcopyrite, Cu-pentlandite, sugakiite [Cu(Fe,Ni)8S8], samaniite [Cu2(Fe,Ni)7S8], chal-cocite, bornite and cubanite; Table 2), native copper and copper–iron–nickel alloys (Table 1; Fig. 3, 4, 5). Within the partly serpentinized peridotites, the sulfide minerals occur mainly in the serpentine groundmass or in serpentine veins, and only rarely within pyroxene, suggesting that most sulfides formed as secondary phases. In some extensively serpentinized samples, sulfides are finely dispersed in the serpentine groundmass as grains <2 μm and could not be determined microscopically or by EMP analyses.

Pentlandite is the most abundant sulfide mineral and was detected in all of the analyzed samples. The most com-mon association is pentlandite + magnetite + awaruite with traces of pyrrhotite and heazlewoodite. In these sam-ples, magnetite predominantly occurs within the cleavage planes of the pentlandite and with awaruite either as thin veins or as thin rims along the pentlandite grains (Fig. 3a, b). Cobalt-pentlandite was only detected in one sample (SE10_06; Table 1) that is almost entirely serpentinized, with Co-pentlandite being intergrown with awaruite. Pyr-rhotite occurs in a few samples as inclusions or along frac-tures within pentlandite, but was only detected in asso-ciation with Cu-bearing sulfides. Additionally, element distribution maps suggest the presence of pyrrhotite as a reaction rim between pentlandite and magnetite (Fig. 4c, d). Heazlewoodite was observed in two samples (SE10_02, SE_P4; Table 1): (1) as an intergrowth with a decomposed

pentlandite crystal and native copper, located within an altered pyroxene, and (2) as intergrowth with Cu-bearing sulfide and Cu–Fe–Ni alloys within the serpentine mesh texture.

One completely serpentinized sample (SE_P5; Table 1) contains the assemblage pentlandite + violarite + smy-thite + chalcopyrite + magnetite. Violarite and smythite form a sub-microscopic intergrowth, are often partly rimmed by magnetite and contain inclusions of pentlandite (suggesting that violarite and smythite replace pentlandite). Chalcopyrite forms veins and is occasionally also rimmed by magnetite.

Cu-bearing sulfide minerals are present in almost all of the analyzed samples and are located together with the Cu-bearing metals within serpentine veins or the serpen-tine mesh texture, and only rarely occur within pyroxene. Within single-sulfide grains, Cu contents increase from the edge inward, with native copper forming a dendritic rim and the Cu-richest sulfur phases forming either along the edges, within fractures increasingly replacing pentlan-dite or as grains together with Cu-rich pentlandite. The most abundant Cu-phases are Cu-pentlandite (Cu content <3 wt%), sugakiite and samaniite, and the latter two are the Cu-rich pentlandite varieties (Table 2). In rare Cu-rich sulfide grains, the Cu-rich pentlandite is accompanied by cubanite, chalcocite and fine-grained intergrowths of bor-nite and chalcocite, while native copper forms dendritic structures into the serpentine groundmass (Fig. 4a, b). In several samples (SE10_01, SE10_02, SE10_05, SE10_19, SE_P4; Table 1), Cu–Fe–Ni alloys occur as thin veins within or along the edge of Cu-bearing sulfides, and in a few samples, Cu–Fe–Ni alloys occur in grains associated with awaruite. Occasionally, serpentine adjacent to sulfides is enriched in Cu, forming bluish halos (under reflected light) along the sulfide grains and in contact with Cu-enriched areas of sulfides (e.g., Fig. 3c). Pentlandite that is rimmed by magnetite is never intergrown with a Cu-rich phase.

Magnetite occurs in most samples as intergrowths with pentlandite (as described above), but also as grains <2 μm in the center of serpentine veins, as fine-grained veins cut-ting the mesh texture, and in some entirely serpentinized samples overgrowing the serpentine mesh texture as large, euhedral grains (up to 0.3 mm). In these latter samples, magnetite does not occur within serpentine veins. Addition-ally, the serpentine in the groundmass has relatively high Mg# of 93–96 implying that most of the Fe in the system has partitioned into magnetite.

Mineral chemistry

The chemical composition of all of the analyzed sulfide and metal phases is reported in the supplementary material


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Tabl

e 1

Min

eral

ogic

al d

escr

iptio

n of

the

Sant

a E

lena

per

idot

ites

a Lith

olog

y: lz

= lh

erzo

lite;

du

= d

unite

; hz

= h

arzb

urgi

teb O

ccur

renc

e of

opa

que

min

eral

s as

det

erm

ined

by

refle

cted

ligh

t mic

rosc

opy

and

EM

P an

alys

es

Sam

ple

nam

eO

rigi

nal

sam

ple

nam

eC

oord

inat

esL

ithol

ogya

Deg

ree

of

serp

entin

i-za

tion

(%)

Opa

que

min

eral

ogyb

NW

ptl

Co–

ptl

pohz

sug

sam

ccbn

cub

smvl

cpm

gtaw

/Fe–

Ni

allo

ysC

u–Fe

–Ni

allo

ysN

ativ

e C

u

SE10

_01

SE10

_01

10°5

5.58

2′85

°51.

922′

lz30

–40

××

××

××

×SE

10_0

2SE

10_0

210

°55.

582′

85°5

1.92

2′lz

40–5

0×

××

××

××

××

SE10

_05

SE10

_05

10°5

4.93

6′85

°48.

608′

lz30

××

××

××

××

SE10

_06

SE10

_06

10°5

4.93

6′85

°48.

608′

du90

–95

××

××

×SE

10_0

9SE

10_0

910

°54.

325′

85°4

7.16

7′hz

60×

××

××

SE10

_12

SE10

_12

10°5

0.52

4′85

°47.

257′

hz30

–40

××

××

××

SE10

_16

SE10

_16

10°5

0.43

6′85

°45.

178′

lz50

–60

××

××

××

××

SE10

_19

SE10

_19

10°5

1.88

2′85

°40.

408′

lz40

–50

××

××

××

××

×SE

_P1

SE_0

1051

0_3

10°5

0.14

2′85

°47.

560′

du85

–90

Sulfi

des

<2

μm

dis

sem

inat

ed in

the

grou

ndm

ass

coul

d no

t be

dete

rmin

ed b

y E

MP

SE_P

2SE

_010

510_

410

°50.

142′

85°4

7.56

0′du

90–9

5×

×SE

_P3

SE_0

2061

1_4

10°5

4.62

0′85

°54.

397′

hz10

0

SE_P

4SE

_050

111_

610

°51.

580′

85°4

1.08

8′hz

40–5

0×

××

××

×SE

_P5

SE_0

5061

1_19

10°5

4.51

5′85

°54.

821′

hz10

0×

××

××


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(Table S1) and plotted in Fig. 2. The stoichiometric formula of the sulfides and metals mentioned in the text is given in Table 2. Several analyses plot between stoichiometric min-eral compositions suggesting that they represent a mixture of at least two mineral phases that form a sub-microscopic intergrowth. Zn contents in all analyzed sulfides and alloys are <0.7wt%.

Pentlandite

Pentlandite analyses reveal a large range in composi-tion with Fe contents of 10.0–44.2 wt% and Ni contents of 21.0–52.1 wt%, the stoichiometric composition rang-ing from (Fe6.1Ni2.9)S8 to (Fe3.6Ni5.4)S8. Highest Fe con-tents are found in grains that are intergrown with mag-netite ± awaruite. In a few samples, two groups can be characterized as having either low or high Fe contents. Co contents are generally <1 wt% but can be as high as 9.5 wt%; in one sample, it reaches 22.2 wt%. Element dis-tribution maps of several grains further suggest that Co is preferentially incorporated into magnetite rather than pent-landite. Cu contents are low (<0.9 wt%) in most pentland-ites that have been unaffected by Cu-alteration.

Pyrrhotite

All of the analyzed pyrrhotite has elevated Ni contents (3.3–6.0 wt%). Many analyses indicate that pyrrhotite forms a sub-micrometric intergrowth with another phase (e.g., with pentlandite, Cu-pentlandite or Fe–Ni alloy; Fig. 2) and are not single mineral grain analyses. Cu con-tents are <0.7 wt% (except for two analyses), and Co con-tents are <0.4 wt%.

Heazlewoodite

All heazlewoodite analyses indicate slightly elevated Fe contents (1.6–4.2 wt%) with a stoichiometric composi-tion of approximately (Ni2.85Fe0.15)S2. Highest Fe contents are observed in Co-rich heazlewoodite (<3.5–7.5 wt%). All other heazlewoodite analyses yielded Co contents <0.5 wt%.

Violarite and smythite

Analyses of both violarite and smythite yielded low totals, while spot analyses with the EDS indicate the presence of small amounts of oxygen. Smythite has a stoichiometric composition of (Fe4.3Ni4.7)S11 to (Fe3.0Ni6.0)S11 and Co con-tents of up to 3.0 wt%. Violarite has a stoichiometric compo-sition of approximately Fe1.0Ni1.9S4, variable Cu contents (up to 4.4 wt%) and Co contents of <1.3 wt%. The presence of violarite and smythite together with chalcopyrite in one sam-ple (SE_P5; Table 1) will not be discussed in further detail as violarite and smythite are most likely the result of late, low-temperature weathering of pentlandite and pyrrhotite, respectively (Craig 1971; Furukawa and Barnes 1996). Indi-cation for weathering is also given by the detection of oxy-gen in several grains, possibly present as a hydrous phase. Thus, we infer that this sulfide assemblage is unrelated to the formation of native copper discussed in this manuscript.

0

20

40

60

80

100

100

80 60 40 20 0 100

80

60

40

20

0

Fe

S

Ni+Cu+Co

pentlandite

pyritechalcopyrite

chalcocitebornite

awaruite (Ni2Fe)

heazlewooditepyrrhotite

awaruite (Ni3Fe)

Fig. 2 Chemical composition of the analyzed sulfides of the studied peridotites. Filled gray circles indicate elevated Cu contents (>3wt%)

Table 2 Mineral abbreviations and formulas

Mineral name Abbreviation Formula

Awaruite aw Ni2Fe to Ni3Fe

Bornite bn Cu5FeS4

Chalcocite cc Cu2S

Chalcopyrite ccp CuFeS2

Covellite cv CuS

Cubanite cub CuFe2S3

Cuprite cpr Cu2O

Heazlewoodite hz Ni3S2

Magnetite mgt Fe3O4

Pentlandite ptl (Fe,Ni)9S8

Polydymite pd Ni3S4

Pyrite py FeS2

Pyrrhotite po FeS

Samaniite sam Cu2(Fe,Ni)7S8

Smythite sm (Fe,Ni)13S16 to (Fe,Ni)9S11

Sugakiite sug Cu(Fe,Ni)8S8

Taenite ta γ Fe–Ni alloy

Violarite vl FeNi2S4


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Fe–Ni alloys

Analyzed Fe–Ni alloys have a large range in composi-tion from Ni57Fe43 to Ni75Fe25 (Figs. 2, 5b), and only a few measurements could be classified as awaruite (Ni2Fe to Ni3Fe). For the alloys with no significant Cu contents (<1 wt%), this range corresponds to a Ni content of 49.6–72.3 wt% and Fe contents of 21.8–39.6 wt%. Co contents are <2.0 wt%, and S contents are generally <1 wt%.

Cu‑phases

Element mapping of sulfide phases has shown that Cu replaces pentlandite from the edge inward, forming Cu-pentlandite, sugakiite and samaniite, locally bornite and chalcocite and native copper along the edge of the sulfide

grain (Fig. 4). Sugakiite and samaniite are the Cu-bearing pentlandite varieties with Cu occupying one and two sites, respectively, in the crystal structure (Table 2). Stoichiomet-ric compositions are approximately Cu0.9(Fe4.1Ni4.0)S8 to Cu2.5(Fe4.5Ni2.0)S8. Thus, reaction of Cu with pentlandite here resulted in a mixing line between Cu-free pentlandite and samaniite with highest Cu contents of 19.3 wt%, and Cu variably replacing either Fe or Ni in the crystal structure (Co <0.9 wt%). Chalcocite is rare and has slightly elevated Fe and Ni contents (Fe = 1.7–5.1 wt%; Ni = 0.5–2.2 wt%) suggesting a sub-microscopic intergrowth with bornite (Fig. 5a). Similarly, a few measurements are situated on a mixing line between chalcocite and pentlandite. Cubanite was detected in only one sample and has a stoichiometric composition of Cu(Fe1.9Ni0.1)S3. Chalcopyrite was only detected associated with pentlandite, violarite and smythite

Fig. 3 Reflected light images of sulfides from four different samples. a Large pentlandite grain intergrown with magnetite and awaruite (Ni2Fe; sample SE10_19). One EMP analysis suggests the presence of a Fe–Ni–Cu alloy with strongly varying composition. b Pentlan-dite with an Fe–Ni alloy forming thin veins in the pentlandite and magnetite as exsolutions in pentlandite (sample SE10_05). EMP

analyses suggest the presence of rare pyrrhotite as inclusions in pent-landite. c Pentlandite intergrown with Cu-rich pentlandite and native Cu. Blue areas are serpentine with traces of Cu (sample SE10_02). d Pentlandite (~100 μm) intergrown with magnetite, and native Cu and an Fe–Ni alloy along the rim (sample SE10_09)


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in one sample and has low Ni contents (<1.2 wt%) and no detectable Co or Zn. Native copper generally contains strongly variable contents of Fe, Ni and S (Fig. 5b), with S contents up to 10 wt% (Fig. 5a) that may be an analytical artifact due to the fine-grained mineral structure and small grain size of the native copper present in these samples.

Bulk rock chemistry

We determined bulk rock major and trace element composi-tions of twelve samples (Table 3; Fig. 6). For many incom-patible trace elements, concentrations are at or below detec-tion limits, reflecting their low abundance in the whole rock peridotites (e.g., most light rare earth elements (LREE) are below detection limits and are thus not reported in Table 3).

Cu contents are between 3.9 and 80.6 ppm (Table 3; Fig. 6). Overall, samples in which no Cu-bearing mineral assem-blages could be detected microscopically have Cu contents <15 ppm. In contrast, samples that contain abundant Cu-bear-ing sulfides and native copper as determined both by miner-alogical observations and by EMP analyses have Cu contents of 16.8 to 80.6 ppm. In general, the lherzolite samples have slightly higher Cu contents than harzburgites (Table 3). There is no distinct relationship between major or trace element bulk rock composition and Cu content with the exception of Zn, which shows a linear correlation (r2 = 0.86). Cu and Zn contents are within the range of partly serpentinized perido-tites from other locations, but lower than most serpentinites from the Logatchev and the Rainbow hydrothermal systems (Fig. 6a). Compared to a depleted MORB mantle (DMM),

Fig. 4 a Sulfide grain under reflected light (sample SE10_19_E). b Elemental map (individual element maps of Fe, Ni, S and Cu overlain to reveal color differences) of the grain shown in a. The map reveals a dendritic rim of native Cu forming around the pentlandite (ptl) grain, while native Cu also forms in the serpentine groundmass and within the pentlandite as veins. Darker red is either a fine-grained mixture between pentlandite + chalcocite (cc) or sugakiite (sug) + samaniite

(sam). c Sulfide grain under reflected light (sample SE10_02_D). d Elemental map (individual element maps of Fe, Ni, S and Cu overlain to reveal color differences) of the grain shown in c. The map reveals a dendritic Cu rim on one side, the presence of samaniite (sam) replac-ing pentlandite and magnetite exsolution within pentlandite with the possible presence of pyrrhotite at the contact between pentlandite and magnetite


1 3

Page 9 of 21 1079

the typically fluid mobile elements Ba (Fig. 6b), K and Pb show distinct enrichments, with Pb showing up to 100 times DMM values (Fig. 6c). However, the fluid mobile element Sr is not enriched in the Santa Elena peridotites compared to DMM, showing very low concentrations (Fig. 6d). The LREE are strongly fractionated relative to the HREE (Table 3) as is expected from depleted upper mantle peridotites (e.g., Salt-ers and Stracke 2004) and did not significantly suffer high-temperature alteration, which typically enriches the LREE (Boschi et al. 2006b; Paulick et al. 2006).

Discussion

The presence of native copper in ultramafic sequences has previously been ascribed to either a magmatic or a meta-somatic origin. In the samples from the Santa Elena Ophi-olite, the presence of native copper and a diverse sulfide mineralogy is of particular interest in understanding the redox as well as the sulfur conditions during the hydrother-mal alteration (serpentinization) of these peridotites. Native copper of primary magmatic origin generally occurs as sin-gle crystals or globule inclusions within olivine, magnet-ite or chromite phenocrysts and is attributed to formation at very low sulfur fugacities or from sulfur-undersaturated magmas (Barkov et al. 1998; Zhang et al. 2006). As a sec-ondary origin, many authors have attributed native copper and Cu–Fe sulfides in serpentinized ultramafic rocks to low S and O2 fugacities during serpentinization by a strongly reducing fluid (Eckstrand 1975; Frost 1985; Lorand 1987). Alternatively, native copper can be related to the addition of Cu by a hydrothermal fluid, for example in the basement of a peridotite-hosted hydrothermal field.

In the following, we will first discuss the character and impact of metasomatizing fluids on the studied peridotite samples and then discuss two models for the origin of the native copper and Cu-rich sulfide assemblages in order to determine the alteration history of the Santa Elena perido-tites and to gain insight into the hydrothermal processes that occur beneath oceanic spreading centers.

Mineralogical record of highly reducing conditions during serpentinization

Serpentinization leads to some of the most reducing con-ditions in natural systems. During reaction of water with olivine (Eq. 1a, 1b) and pyroxene (Eq. 2), Fe2+ in the pri-mary mineral phases is oxidized to Fe3+ with water as the reducing agent forming H2 according to the following set of reactions (Bach et al. 2006; Beard et al. 2009; Frost et al. 2013; Schwarzenbach et al. 2013a):

In oceanic serpentinites, hydrogen fugacities tend to decrease with progressive serpentinization as the rock becomes depleted in olivine, which is typically associ-ated with increased fracturing allowing more seawater to enter the system, thus, additionally increasing water–rock

(1a)Olivine + fluid = Mg-rich serpentine + Fe-rich brucite

(1b)

Mg-rich serpentine + Fe-rich brucite + fluid

= Mg-rich serpentine + Mg-rich brucite

+ magnetite + H2

(2)Pyroxene + fluid = serpentine + talc

covellitechalcocite

0

20

40

60

80

100

100

80 60 40 20 0

100

80

60

40

20

0

bornite

chalcopyritecubanite

samaniite

sugakiite

Cu

S

Fe+Ni

(a)

Fe

Ni

20

40

60

80

100

100

80 60 40 20 0

100

80

60

40

20

0

0

awaruite (Ni2Fe) awaruite (Ni3Fe)

Cu

Ni

(b)

Fig. 5 a Composition of the Cu sulfides (includes analyses of sulfides with Cu contents >1 wt%) in S–Cu–Fe + Ni space. b Fe–Ni–Cu alloys in the serpentinized peridotites plotted in Cu–Fe–Ni space. All plotted analyses have S contents of <3 wt%. The analyzed sam-ples plot on a mixing trend between a Fe–Ni end-member that has a composition similar to awaruite and native Cu


1 3

1079 Page 10 of 21

Tabl

e 3

Maj

or (

in w

t%)

and

trac

e el

emen

t com

posi

tions

(in

ppm

) of

the

Sant

a E

lena

per

idot

ites

<l.o

.d =

bel

ow li

mit

of d

etec

tion

Sam

ple

SE10

_02

SE10

_05

SE10

_06

SE10

_09

SE10

_12

SE10

_16

SE10

_19

SE_P

1SE

_P2

SE_P

3SE

_P4

SE_P

5R

ange

of

dete

ctio

n lim

it

Maj

or e

lem

ents

(in

wt%

)

SiO

243

.80

43.9

238

.01

42.1

243

.36

42.4

343

.44

39.5

939

.63

42.6

842

.82

43.0

5<

0.01

TiO

20.

060.

05<

l.o.d

0.01

0.04

0.04

0.05

<l.o

.d<

l.o.d

<l.o

.d0.

020.

01<

0.01

Al 2

O3

3.14

2.65

0.13

1.11

1.99

2.04

2.50

0.29

0.45

0.40

1.21

0.72

<0.

01

Fe 2

O3

8.92

8.56

8.24

8.40

8.69

8.48

8.36

7.28

7.62

8.79

8.67

8.89

<0.

01

MnO

0.14

0.14

0.12

0.13

0.14

0.13

0.13

0.11

0.11

0.11

0.14

0.13

<0.

01

MgO

38.4

039

.92

44.7

242

.25

40.8

039

.58

38.7

846

.91

46.6

040

.55

42.9

040

.27

<0.

01

CaO

2.71

2.16

0.14

1.03

1.71

1.86

2.21

0.09

0.04

0.02

0.82

0.05

<0.

01

K2O

0.03

0.03

0.03

0.03

0.03

0.02

0.03

0.03

0.03

0.03

0.03

0.03

<0.

01

P2O

50.

020.

02<

l.o.d

<l.o

.d<

l.o.d

0.01

0.01

<l.o

.d<

l.o.d

<l.o

.d<

l.o.d

0.01

<0.

01

Tra

ce e

lem

ents

(in

ppm

)

Ni

1,64

41,

666

2,07

018

181,

726

1,71

01,

598

2,67

32,

326

2,38

21,

833

2,01

76–

9

Sc

23.9

23.4

19.1

20.5

20.9

22.1

21.1

24.6

23.1

24.2

23.6

22.7

<2.

2

V66

.957

.513

.433

.747

.550

.660

.29.

311

.414

.035

.937

.0<

0.5

Cr

2,50

0.6

2,62

0.3

2,68

0.6

2,47

4.9

2,24

9.6

2,26

7.1

2,34

4.5

2,42

4.7

3,04

5.7

3,34

7.6

2,95

8.2

3,12

1.5

<1.

4

Co

102.

710

1.5

116.

910

9.3

105.

310

2.5

97.5

120.

412

3.1

126.

111

0.4

117.

1<

2.2

Cu

32.4

25.1

6.5

16.8

20.1

80.8

26.3

29.4

68.2

3.9

9.2

24.2

3.8–

5.4

Zn

55.3

49.1

45.8

47.3

46.3

68.7

46.9

46.3

55.2

42.1

51.3

49.5

2.2–

2.9

Sr

1.04

1.19

1.64

0.21

0.18

0.32

0.31

2.07

1.42

1.09

0.83

1.43

0.06

–0.0

9

Y2.

452.

04<

l.o.d

.0.

541.

481.

842.

23<

l.o.d

.<

l.o.d

.0.

090.

500.

110.

08–0

.11

Zr

0.35

0.51

0.33

<l.o

.d.

0.29

0.29

0.54

<l.o

.d.

<l.o

.d.

0.21

0.17

<l.o

.d.

0.12

–0.2

2

Sn

<l.o

.d.

0.64

0.88

0.26

1.05

0.33

0.37

0.33

0.28

0.91

0.31

0.32

0.18

–0.2

6

Ba

1.78

1.66

2.00

1.44

1.46

1.91

1.60

4.69

5.04

1.88

1.58

1.49

0.02

–0.0

4

La

<l.o

.d.

<l.o

.d.

0.07

<l.o

.d.

<l.o

.d.

<l.o

.d.

<l.o

.d.

<l.o

.d.

<l.o

.d.

0.05

<l.o

.d.

<l.o

.d.

0.02

–0.0

5

Ce

0.03

<l.o

.d.

0.06

<l.o

.d.

<l.o

.d.

<l.o

.d.

0.03

<l.o

.d.

<l.o

.d.

0.06

<l.o

.d.

<l.o

.d.

0.02

Dy

0.40

0.29

<l.o

.d.

<l.o

.d.

0.23

0.32

0.35

<l.o

.d.

<l.o

.d.

<l.o

.d.

<l.o

.d.

<l.o

.d.

0.09

–0.1

6

Ho

0.09

0.08

<l.o

.d.

0.02

0.06

0.07

0.09

<l.o

.d.

<l.o

.d.

<l.o

.d.

0.02

<l.o

.d.

0.01

–0.0

3

Er

0.26

0.23

<l.o

.d.

0.05

0.18

0.20

0.28

<l.o

.d.

<l.o

.d.

<l.o

.d.

0.06

<l.o

.d.

0.02

–0.0

7

Yb

0.31

0.23

<l.o

.d.

0.07

0.22

0.23

0.30

<l.o

.d.

<l.o

.d.

<l.o

.d.

0.07

<l.o

.d.

0.04

–0.0

9

Pb

0.39

0.21

0.42

<l.o

.d.

0.32

17.6

6<

l.o.d

.1.

431.

50<

l.o.d

.0.

280.

180.

16–0

.21


1 3

Page 11 of 21 1079

ratios and oxygen fugacities (Frost 1985; Alt and Shanks 1998; Delacour et al. 2008c; Schwarzenbach et al. 2012). Oxygen fugacity is effectively fixed until olivine is no longer in contact with fluid (Frost 1985; Schwarzenbach et al. 2013a). As the opaque mineralogy is a function of hydrogen/oxygen and sulfur fugacities (Eckstrand 1975; Frost 1985; Klein and Bach 2009), entirely serpentinized peridotites usually preserve an opaque mineral assem-blage consisting of high-sulfur assemblages (e.g., pyrite, vaesite) and magnetite and/or hematite (Delacour et al. 2008b; Schwarzenbach et al. 2012). In contrast, high H2 fugacities typical for the initial stages of serpentiniza-tion allow the stability of native metals and Fe–Ni alloys (Frost 1985; Alt and Shanks 1998). Accordingly, many peridotite-hosted hydrothermal systems record a typical redox gradient with highly reducing conditions during

initial serpentinization trending to more oxidizing con-ditions during late stages of serpentinization (Alt and Shanks 1998; Delacour et al. 2008b; Schwarzenbach et al. 2012).

In the Santa Elena peridotites, strongly reducing con-ditions during serpentinization are recorded by the pres-ence of awaruite and rare heazlewoodite. The assemblage pentlandite + awaruite + magnetite is formed as the result of the destabilization of pentlandite and is commonly observed in serpentinized peridotites (Eckstrand 1975; Per-etti et al. 1992; Klein and Bach 2009). This desulfurization of pentlandite can be described by the reaction (Eq. 3, after Klein and Bach 2009):

(3)

Ni4.5Fe4.5S8 + 4H2 + 4H2O = 1.5Ni3Fe + Fe3O4 + 8H2S

Ba0.01 0.1 1 10 100 1000

0.1

1

10

100

1000

10000

Cu

(b)

0.1

1

10

100

1000

10000

Cu

Pb0001100.0 1001010.01 0.1

(c)

0.1

1

10

100

1000

10000

Cu

Sr000111.010.0 10 100 10000

(d)

0.1

1

10

100

1000

10000

Cu

Zn10 100 1000

Santa ElenaLogatchev serpentinitesRainbow serpentinitesMariana forearc15° 20’ fracture zoneAtlantis MassifDMM

10000

(a)

Fig. 6 Bulk rock compositions of the Santa Elena peridotites: Cu variations shown against a Zn, b Ba, c Pb and d Sr contents, and compared with ultramafic rocks from other locations. Additional data are serpentinites from the basement of the Logatchev hydrothermal field (Augustin et al. 2012), the basement of the Rainbow hydrother-mal field (Marques et al. 2007), the Mid-Atlantic Ridge 15°20′N frac-ture zone (ODP Leg 209) (Paulick et al. 2006; Kodolanyi et al. 2012),

the Atlantis Massif (Delacour et al. 2008c) and the Mariana fore arc conical seamounts (ODP Leg 125) (Savov et al. 2005; Kodolanyi et al. 2012). This shows that the Santa Elena peridotites have Cu, Zn and Ba contents within the range of ultramafic rocks from similar set-tings and are at the high end of Pb content and low end of Sr content of such rocks (see text for “Discussion”). DMM (depleted MORB mantle) after Salters and Stracke (2004)


1 3

1079 Page 12 of 21

Additionally, the presence of awaruite in almost all of the samples, even including the entirely serpentinized peri-dotites, suggests low water–rock ratios. Assuming serpen-tinization temperatures of <250 °C (as will be discussed in the next section), stabilization of awaruite requires water–rock ratios of <1 during water–rock interaction (approxi-mated after calculations by Klein et al. 2009). Such low water–rock ratios could have maintained prolonged highly reducing conditions even in almost entirely serpentinized samples during alteration up to the complete replacement of the primary minerals by serpentine in some samples.

Proxies for alteration temperatures

Serpentinization of ultramafic rocks in oceanic settings generally takes place between 150 and 500 °C. Tempera-tures are largely controlled by the tectonic setting and the presence or absence of a magmatic heat source, and may vary over the time span of serpentinization (e.g., Cannat et al. 1992; Früh-Green et al. 1996; Agrinier and Cannat 1997; Schwarzenbach et al. 2013b). The upper tempera-ture limit of mineral alteration is controlled by the stability of olivine and pyroxene. In the presence of a fluid and at elevated SiO2 activities, pyroxene decomposition is faster at >350–400 °C than decomposition of olivine, while at <250 °C pyroxene is more stable and serpentinization of olivine is faster than breakdown of pyroxene (Martin and Fyfe 1970; Bach et al. 2004; Frost and Beard 2007). In the Santa Elena peridotites, pyroxene is almost intact, while

olivine is strongly serpentinized. This relationship suggests serpentinization temperatures of <250 °C.

Temperatures can also be estimated from Fe–Ni–S phase assemblages (Fig. 7a), using, for example, the composi-tion of co-existing pentlandite and awaruite, or Ni contents in pyrrhotite (Craig 1973; Misra and Fleet 1973). In three samples, coexisting pentlandite and awaruite suggest tem-peratures of around 250 °C (Fig. 7a; Craig 1973; Vaughan and Craig 1997). However, several pentlandite + awaruite assemblages suggest higher formation temperatures of 300–400 °C (after Craig 1973; Vaughan and Craig 1997). Similarly, elevated Ni contents (3.3–6.0 wt%) in pyr-rhotite detected in a few samples suggest temperatures ≥400 °C, where complete pyrrhotite–millerite monosulfide solid solution occurs (Craig 1973; Misra and Fleet 1973). Another temperature proxy is the pentlandite composi-tion in the ternary system Fe9S8–Ni9S8–Co9S8 suggested by Kaneda et al. (1986) (Fig. 7b). Most of the analyzed pentlandite plots within the entire field that is stable down to temperatures of <200 °C. However, pentlandite com-positions in one sample preserve formation temperatures of >300 °C and possibly even >600 °C (Fig. 7b). In sum-mary, these temperature proxies suggest multiple phases of alteration with temperatures around 200 °C likely reflecting the main stage of serpentinization and a higher temperature fluid influx event at 350–400 °C, which we infer to postdate serpentinization (see below). We infer that even higher tem-peratures (>400 °C) are associated with primary pentlan-dite and pyrrhotite still partially preserved in the samples.

NiS2FeS2

Ni3S4

NiS

Ni3S2

Fe7S8FeS

FeNi2S4

(Fe,Ni)9S8

Mss

iNeF

S

FeNi3±x

)) SS(Fe Ni)F(F

MsssMs

i)))

222S

(F(F(FF

ssFF

)i)99

seN

SE10_19SE10_12SE10_01

600°C

500°C

400°C

300°C

200°C

0

20

40

60

80

100

100

80 60 40 20 0 100

80

60

40

20

0

Co9S8

Fe9S8 Ni9S8

(b)(a)

Fig. 7 a Mineral stabilities in the ternary diagram Fe–Ni–S with coexisting pentlandite and awaruite compositions from three different samples. Phase relations are for 250 °C, redrawn from Craig (1973). b Pentlandite analyses in the ternary system Fe9S8–Ni9S8–Co9S8,

with temperature-dependent stability fields for the formation of pent-landite (after Kaneda et al. 1986). Pentlandite occupies the widest range in composition at 500 °C, separating into two stability fields at temperatures <200 °C


1 3

Page 13 of 21 1079

Influence of changing redox conditions and the formation of native copper

Cu-bearing sulfide mineral assemblages and native cop-per have been described in several ultramafic sequences, revealing both similarities and differences to the obser-vations presented in this study. In the Zambales ophiolite (Philippines), Abrajano and Pasteris (1989) describe a pri-mary magmatic assemblage consisting of pyrrhotite, pent-landite, chalcopyrite, magnetite and a secondary assem-blage consisting of chalcocite, cubanite, digenite, bornite and idaite together with native copper. They attribute the secondary assemblage to re-equilibration and alteration of Cu-bearing sulfides as a result of reducing conditions dur-ing serpentinization. Similarly, Lorand (1987) suggest that native copper in lherzolites and harzburgites from the Bay of Islands Ophiolite (Newfoundland) formed as an altera-tion product and through breakdown of primary sulfides due to interaction with strongly reducing fluids during serpentinization. The peridotites from the Bay of Islands Ophiolite preserve similar textures to those observed in the Santa Elena peridotites with native copper or awaruite rims around pentlandite. However, native copper is sometimes located within serpentine veinlets and Lorand (1987) did not report the occurrence of samaniite, sugakiite and sub-micrometric intergrowths of various sulfide minerals. Cu-bearing mineral assemblages have also been described in plagioclase lherzolites from the Horoman peridotite com-plex (Japan; Kitakaze 2008; Kitakaze et al. 2011), where samaniite and sugakiite were described in detail for the first time. In these peridotites, native copper is attributed to a magmatic origin with formation through crystallization of an immiscible metallic liquid. The primary copper was sub-sequently partly oxidized and formed a secondary Cu-rich mineral assemblage (Ikehata and Hirata 2012).

Most of the sulfides observed in the Santa Elena peri-dotites are located within serpentine veins and suggest a secondary formation for the sulfide and metal mineral assemblages. Only a few pentlandite grains located within ortho- and clinopyroxene suggest a primary, magmatic origin, even though several of these sulfide grains were affected by the metasomatizing fluid while the pyroxene was partly serpentinized. As discussed above, the most common association of pentlandite + awaruite + magnet-ite can be attributed to formation during serpentinization. As proposed for the Zambales and Bays of Islands ophi-olite, highly reducing conditions caused destabilization of pentlandite and the re-equilibration of the primary sulfide assemblage (Klein and Bach 2009). The mineralogical observations and the element distribution maps suggest that the Cu-bearing mineral assemblages formed as a secondary feature after pentlandite re-equilibration. Evidence thereof is that awaruite is partly replaced by Cu–Fe–Ni alloys and

that pentlandite–awaruite intergrowths that are rimmed by magnetite, which likely formed as a result of desulfuriza-tion of pentlandite, are never intergrown with Cu-bearing phases. Hence, the Cu-bearing phases most likely formed as fluid conditions changed (chemically and possibly also to higher temperatures) subsequent to the main stage of ser-pentinization. This locally overprinted the opaque mineral assemblages to form the Cu-bearing mineral assemblages. Evidence that several different fluids interacted with the peridotites is suggested by the preservation of indicators of variable alteration temperatures, as described above.

Stability of native copper as a function of oxygen and sulfur fugacity

To test for the influence of changing temperature and oxy-gen and sulfur fugacities on the sulfide assemblages, we calculated equilibrium mineral stability fields in the system MgO, SiO2, Fe, S, Cu, H and O at 0.5 kbars and tempera-tures of 200 and 350 °C (Fig. 8). Gibbs free energy mini-mization with Perple_X (Connolly 2005) used thermo-dynamic end-member data for silicate, oxide and sulfide phases from supcrt92 (Johnson et al. 1992 and references therein). Olivine and orthopyroxene were modeled as binary Fe–Mg solutions, and the primary bulk rock compo-sition was modeled as a system comprising of 66.6 mol% olivine and 33.3 mol% orthopyroxene, representing approximately the composition of the studied harzburgites (Table 3). Copper was added to form native copper and Cu sulfides, and Fe was added as a component to permit stabil-ity of iron sulfides, native metal, FeO in silicates and Fe2O3 in oxides. Consideration of MgO (rather than Mg) assumes that all Mg is divalent, unlike the case for Fe. We did not include Ni and, therefore, pentlandite, due to the lack of thermodynamic data for Cu-bearing pentlandite, sugakiite and samaniite, which are the most abundant Cu-bearing sulfide phases in the studied samples, and the lack of well-constrained mixing models for Cu in pentlandite. We thus assume that the stability fields of the Cu-bearing phases in the calculations can be used to approximate the changing conditions observed in the samples and that the addition of Ni will have limited effect on the Fe–Cu assemblages.

As a function of oxygen and sulfur fugacities, native copper has a relatively large stability field at both 200 and 350 °C (shaded area in Fig. 8a, b). At very reducing condi-tions and low sulfur fugacities, the stability field of native copper overlaps with the stability field of native iron and at fairly oxidizing conditions it partly overlaps with hem-atite (Fig. 8). This implies that changing oxygen fugac-ity has a limited effect on the stability of native copper if sulfur fugacities are low. In contrast, the sulfur fugacity of the fluid has a significant effect on the Cu-bearing mineral assemblages. At low fO2, an increase in fS2 results in the


1 3

1079 Page 14 of 21

formation of bornite followed by formation of chalcopy-rite. At high fO2, an increase in fS2 results in the formation of chalcocite before forming bornite, chalcopyrite, then covalite and higher sulfur phases (Fig. 8a, b). Accordingly, if primary chalcopyrite is present, a decrease in fS2 should first result in the formation of bornite and/or chalcocite and then native copper. Importantly, higher temperatures significantly increase the stability field of native copper to both higher oxygen and sulfur fugacities. Thus, a fluid with a given composition could produce chalcopyrite and pyr-rhotite at 200 °C while producing native copper at 350 °C (Fig. 8a, b). Additionally, H2S and H2O are favored at low temperatures.

In summary, the calculations show that decomposition of chalcopyrite caused by decreasing fO2 and fS2 results in the formation of bornite, chalcocite and, at lowest fS2, native copper. Conversely, they show that Cu-bearing sulfides can form by combined oxidation and sulfidation of primary native copper. However, since all of the stud-ied assemblages point to highly reducing conditions, we exclude the possibility that the Cu-bearing sulfides in the Santa Elena peridotites formed through late stage cou-pled oxidation and sulfidation of primary native copper, as for example suggested for Cu-bearing assemblages in the Horoman Complex (Ikehata and Hirata 2012). Thus, we infer that the observed Cu-bearing assemblages may have formed due to interaction with sulfur-depleted fluids

during serpentinization at <250 °C. This may have caused primary sulfides to (partly) breakdown, releasing Cu into the fluid and forming a new mineral assemblage, e.g., Cu-bearing pentlandite, sugakiite, samaniite, rare bornite and native copper. Note that none of the samples that contain native copper preserve any chalcopyrite (chalcopyrite was only detected in one sample in association with vio-larite and smythite, which is inferred to have formed as a result of late, low-temperature weathering), suggesting that all initial chalcopyrite was replaced during desulfuri-zation. An alternative hypothesis requires incorporation of Cu into pentlandite and subsequent release of this Cu dur-ing pentlandite breakdown. In the next section, we further test whether decomposition of these primary sulfides due to highly reducing and sulfur-depleted serpentinizing con-ditions could have resulted in the formation of the rela-tively large amounts of native copper observed in the thin sections.

A mass balance test of copper exsolution from primary sulfides

A mass balance calculation based on element distribu-tion maps of three representative sulfide grains allows us to evaluate whether breakdown of the Cu component of either (1) pentlandite or (2) pentlandite + chalcopyrite, due to thermodynamic instability during highly reducing

+H2

+H2O

+H2S

Ol Opx Cu Fe

Ol Opx Cu

Ol Cu Tcl

Ol Cu Atg Mag

Brc Cu Atg Mag 2 6

7

9

10

11

+H2O+H2O+H2

+H2O

+H2S

+H2

+H

Ol Opx Cu Fe

Ol Opx Cu

Ol Cu Tcl

Ol Cu Atg Mag

Brc Cu Atg Mag

Brc CuAtg Hem

Brc CprAtg Hem

Brc AtgMag Cc

Brc Atg Cc Hem

OlOpxBnPo

Ol OpxCcp Po

Ol TclCcp Po

Ol O

px P

y B

n

Ol O

px C

v P

y B

rc C

v A

tg H

em

2

3

4

56

7

8 9

10

11

12

13

1: Ol Atg Mag Bn 2: Brc Atg Mag Bn 3: Brc Atg Bn Hem 4: Brc Atg Mag Ccp5: Brc Atg Ccp Po6: Brc Py Atg Ccp7: Brc Py Atg Bn8: Ol Atg Ccp Po9: Ol Py Atg Ccp10: Ol Py Atg Bn11: Ol Py Tcl Ccp12: Ol Py Tcl Bn13: Ol Cv Py Tcl


1

-25 -20 -15 -10 -5

log f S2

log

f O2

-60

-55

-50

-45

-40

-35

-30

-25

-20

+H2

+H2O+H2O+H2+H2

Ol O

px C

u P

o

Ol Cu Tcl Atg 2

5

9

Ol Opx Cu Fe

Ol Opx Cu

Ol Cu Tcl

Ol Cu Atg Mag

Ol Cu Atg Hem

Ol Cpr Atg Hem

Ol O

px C

u P

o

Ol O

px B

n P

o

Ol AtgMag Cc

Ol AtgCc Hem

Ol O

px C

cp P

o

Ol Opx Cu Tcl

Ol Cu Tcl Atg 1 2

3 4 57 8 9

6

1: Ol Atg Mag Bn2: Ol Atg Mag Ccp3: Ol Tcl Cc4: Ol Tcl Bn5: Ol Tcl Ccp6: Ol Tcl Ccp Po7: Ol Opx Cc8: Ol Opx Bn9: Ol Opx Ccp


-25 -20 -15 -10 -5

log f S2lo

g f O

2

-60

-55

-50

-45

-40

-35

-30

-25

-20C ̊053 (b)C ̊002(a)

+H2

+H2O

+H2S

Ol Opx Cu Fe

Ol Opx Cu

Ol Cu Tcl

Ol Cu Atg Mag

Brc Cu Atg Mag 2 6

7

9

10

11

+H2O+H2O+H2

+H2O

+H2S

2S

Ol Opx Cu Fe

Ol Opx Cu

Ol Cu Tcl

Ol Cu Atg Mag

Brc Cu Atg Mag

Brc CuAtg Hem

Brc CprAtg Hem

Brc AtgMag Cc

Brc Atg Cc Hem

OlOpxBnPo

Ol OpxCcp Po

Ol TclCcp Po

Ol O

px P

y B

n

Ol O

px C

v P

y B

rc C

v A

tg H

em

2

3

4

56

7

8 9

10

11

12

13



1

-25 -20 -15 -10 -5

log f S2

log

f O2

-60

-55

-50

-45

-40

-35

-30

-25

-20

+H2

+H2O+H2O+H2+H2

Ol O

px C

u P

o

Ol Cu Tcl Atg 2

5

9

Ol Opx Cu Fe

Ol Opx Cu

Ol Cu Tcl

Ol Cu Atg Mag

Ol Cu Atg Hem

Ol Cpr Atg Hem

Ol O

px C

u P

o

Ol O

px B

n P

o

Ol AtgMag Cc

Ol AtgCc Hem

Ol O

px C

cp P

o

Ol Opx Cu Tcl

Ol Cu Tcl Atg 1 2

3 4 57 8 9

6



-25 -20 -15 -10 -5

log f S2lo

g f O

2

-60

-55

-50

-45

-40

-35

-30

-25

-20C ̊053 (b)C ̊002(a)

Fig. 8 Stabilities of mineral assemblages in the system MgO, SiO2, Fe, S, Cu, H and O calculated by Gibbs free energy minimi-zation. Stability fields are calculated at 0.5 kbars and temperatures of a 200 °C, representing the main stage of serpentinization, and b

350 °C, representing conditions during a higher temperature postser-pentinization event. The shaded areas represent the stability field of native copper. Colored fields represent the volatile phase (H2, H2S or H2O) in equilibrium with the assemblage


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Page 15 of 21 1079

and sulfur-depleted conditions, formed the native copper observed in the samples. EMP point analyses indicate that the Cu rims revealed by element mapping are relatively pure Cu with generally <1 wt% of Fe, Ni and S, while highly variable Cu contents are present within the outline of the original pentlandite (e.g., Fig. 4b, d). We estimated total Cu contents using the area of the Cu rim on the ele-ment maps of three different sulfide grains. In model 1, we compare this value with the amount of Cu present in the primary pentlandite (area calculated using the out-line of the entire pentlandite without the Cu rim)—Cu contents in pentlandite are generally low and <0.4 wt% (Puchelt et al. 1996; Luguet et al. 2003; Schwarzenbach et al. 2012). In model 2, we compare the calculated value with the amount of Cu present in primary pentlandite and variable amounts of primary chalcopyrite, which is often found as an accessory phase in peridotites (Alt et al. 2007; Delacour et al. 2008b; Schwarzenbach et al. 2012). For both calculations, we use the following assumptions: (1) The relative amount of native copper formed upon breakdown of the primary sulfides is equal in two and three dimensions (i.e., the area of the element maps can be used as a proxy for the total volume), (2) the area of the primary sulfides is approximated by the contour of the mapped sulfide and assumes that no fractures were present in the primary grain, (3) the area calculated as native copper contains no traces of any other elements (e.g., Fe, Ni, S) and (4) all of the Cu originally present in the primary sulfides is expelled and forms now native copper. Additionally, we assume for model 1 that the pri-mary pentlandite had a composition of (Fe5.7Ni3.2Cu0.1)S8, which corresponds to 0.83 wt% Cu (0.44 vol% Cu) and is assumed to be a maximum value of Cu in pent-landite. For model 2, we assume that the primary sulfide was comprised of pentlandite and between 2 and 25 vol% chalcopyrite.

Based on the above-mentioned assumptions, we calcu-late for model 1 that the fraction of Cu originally present in the pentlandite was 10–15 times lower than the volume of native copper now observed in the thin sections. For model 2, we calculate that the fraction of Cu originally present in the primary phases was 6–10 times (with 2 % of the sulfide being chalcopyrite), 3–6 times (with 5 % ccp) and <2 times (with 25 % ccp) lower than the vol-ume of native copper now observed. Note that chalcopy-rite is typically only present in trace amounts in perido-tites and is best represented by assuming 2 vol% (or less) of chalcopyrite. Thus, for both models, the calculations suggest either that the initial peridotite was unusually Cu rich or that an external source of additional Cu is neces-sary to produce the native copper abundance observed here. However, reconsidering the above-listed assump-tions, several additional points need to be considered:

(1) during desulfurization, not all of the Cu is expelled from the pentlandite, with up to 0.3 wt% of Cu still pre-sent in the analyzed pentlandite. This suggests that the amount of native copper produced should be even smaller if all Cu originated from sulfides. (2) No chalcopyrite was found in association with native copper, samaniite or sug-akiite. As significant amounts of primary pentlandite are still preserved in the samples, we would expect at least trace amounts of primary chalcopyrite to also still be pre-sent if native copper formed as a result of breakdown of primary chalcopyrite. (3) On the other hand, given that there are also variable amounts of Fe, Ni and S in the Cu rims (especially along the edges), the above calculation may lead to an overestimation of the calculated area for native copper. Assuming that only 50 vol% of the calcu-lated area is pure Cu, decomposition of 90 vol% pent-landite and 10 vol% chalcopyrite could possibly have produced the observed native copper rims. (4) Cu may be mobile in the fluid and could have locally been accu-mulated around sulfide grains. The three calculated areas would therefore not represent the Cu expelled from just the adjacent pentlandite, but the amount expelled from several pentlandite grains. In summary, both points (1) and (2) strongly support that an external source for Cu was necessary, while points (3) and (4) relax this assertion and suggest that native copper could have formed through breakdown of the primary sulfides if chalcopyrite was present and/or if Cu was mobile on a thin section scale.

Evidence for an external source for copper in the serpentinites

The average upper mantle contains little Cu (~25–30 ppm Cu; Sun 1982; Salters and Stracke 2004) with Cu con-tents reported for abyssal peridotites generally below 30–50 ppm (Niu 2004; Savov et al. 2005; Paulick et al. 2006; Delacour et al. 2008c; Zeng et al. 2012) and between 3 and 40 ppm in orogenic peridotites (Garuti et al. 1984; Lorand 1989a and references therein). Overall, Cu con-tents should be lower in harzburgites than in lherzolites as Cu is incompatible during partial melting (Lorand 1989a). Cu contents of the Santa Elena peridotites are between 3.9 and 80.6 ppm and are within the range of variably serpen-tinized peridotites from the Mid-Atlantic Ridge 15°20′N fracture zone (ODP Leg 209), the Mariana forearc coni-cal seamounts (ODP Leg 125) and the Atlantis Massif (IODP Leg 304/305) (Fig. 6). The Cu contents of several harzburgites are similar to those of the lherzolite samples, while some of the dunites have higher Cu contents than the harzburgites. Together this suggests that the Cu con-tents are not simply related to the amount of melt that may have been extracted and that they may represent secondary enrichment.


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Similarities to ultramafic rocks affected by Black Smoker type fluids

Cu-rich sulfide deposits are typically found in correlation with heat-driven seawater circulation reacting with crus-tal to upper mantle rocks, which are both the source of the heat and the metals that eventually form the ore depos-its (e.g., von Damm 1990; German et al. 1993; Rona and Scott 1993; Fouquet et al. 1996; Candela 2003; German and von Damm 2003). Volcanogenic massive sulfide ore deposits (VMS deposits) are generally considered to be their equivalent on the continent and are characterized by high Cu and variably elevated Zn, Pb, Ag and Au contents (e.g., Candela 2003). Active seafloor hydrothermal systems vent hot (300 to >400 °C), acidic and metal-rich fluids and are abundant along ocean ridges; well-studied examples include TAG, Broken Spur and Lucky Strike (German et al. 1993; Langmuir et al. 1997; German and von Damm 2003). These are hosted by mafic rocks and comprise a complex system of hydrothermal fluid circulation in the subsurface that results in the formation of the massif sulfide depos-its. Several hydrothermal systems have been discovered that are hosted by ultramafic rocks, but vent acidic, metal-rich fluids at temperatures >350 °C, while the ultramafic basement is undergoing serpentinization (Fouquet et al. 1997; Charlou et al. 2002; Douville et al. 2002; Allen and Seyfried 2004; Schmidt et al. 2007). The Rainbow and Logatchev hydrothermal systems are two examples of such ultramafic-hosted hydrothermal systems and are both located along the Mid-Atlantic Ridge (Fouquet et al. 1997; Petersen et al. 2009). In both systems, it has been sug-gested that the heat is supplied by gabbroic intrusions in the footwall of the hydrothermal system, where the hydro-thermal fluids leach metals in the mafic rocks (Allen and Seyfried 2004; Petersen et al. 2009), while the serpentinites preserve a low-temperature seawater signature (<250 °C). The systems typically comprise a range from stockwork serpentinites with disseminated sulfides or sulfides in vein-lets to massive Cu sulfides, suggesting that in both hydro-thermal systems, Cu is locally added by hydrothermal flu-ids (leached at high temperatures from mafic sequences) resulting in a heterogeneous element distribution. The ultramafic basement of the Logatchev and Rainbow fields has highly variable Cu contents of 3 ppm to 2.692 wt% for serpentinites at Logatchev (Augustin et al. 2008) and <10 ppm to 542 ppm in serpentinites and up to 5,033 ppm in the stockwork at Rainbow (Marques et al. 2007). In both cases, Cu/Zn ratios are typically <1 (Marques et al. 2007; Augustin et al. 2012).

In the Santa Elena peridotites, Cu-rich sulfides and native copper occur in most samples (see Table 1). Cu con-tents of the bulk rocks are generally lower in samples with little Cu detected in thin section (<15 ppm), while samples

with significant amounts of microscopically obvious Cu sulfides contain up to 81 ppm Cu. Cu/Zn ratios are below 1.2 and average 0.5, similar to the ratios of the Logatchev and Rainbow serpentinites. Most sulfides and all of the Cu-rich sulfides are located within serpentine and clearly indi-cate a secondary origin, suggesting as described above that the Cu-bearing mineral assemblage formed after the main stage of serpentinization. Even though Cu contents are rela-tively low compared to the basement rocks from Logatchev and Rainbow, the peridotites from Santa Elena could rep-resent sections of a similar hydrothermal system, where Cu-bearing, high-temperature hydrothermal fluids, pos-sibly derived from adjacent mafic lithologies, could have locally interacted with the peridotites after serpentinization. Similarly, Marques et al. (2007) describe that localized magmatic/hydrothermal sulfide mineralization postdate serpentinization in the basement of the Rainbow hydro-thermal system, while the basement rocks preserve typi-cal serpentine mesh and hour glass textures. The studies of Logatchev and Rainbow infer that the main stage of serpen-tinization of the ultramafic basement takes place through interaction with seawater, without extensive sulfide min-eralization. Subsequent interaction with high-temperature, acidic hydrothermal fluids occurs primarily in a localized upflow zone, forming stockwork and massif sulfide depos-its (Marques et al. 2006, 2007; Augustin et al. 2012). Away from the upflow zone interaction with high-temperature flu-ids probably only occurs locally, resulting in heterogeneous metal enrichments. Accordingly, the peridotites from Santa Elena may have only locally been affected by high-tem-perature, Cu-bearing hydrothermal fluids, which are likely responsible for the formation of the higher temperature sig-natures (350–400 °C) detected in a few samples.

Evidence from trace elements

Evidence for interaction with high-temperature hydrother-mal fluids is often preserved in trace element compositions, e.g., in flat REE patterns or elevated concentrations of fluid mobile elements (Niu 2004; Boschi et al. 2006a; Paulick et al. 2006). Serpentinites from mid-ocean ridge environ-ments at which substantial melt was generated characteris-tically preserve strongly depleted trace element concentra-tions. As a result of subsequent serpentinization, hydrated peridotites show relative enrichments in the fluid mobile elements (Cl, B, Rb, U, Pb, Sb, Sr and Li), even though concentrations are usually below primitive mantle values (Niu 2004; Kodolanyi et al. 2012; Deschamps et al. 2013). The Santa Elena peridotites generally have very low incom-patible trace element concentrations, with most of the ana-lyzed REEs at or below the detection limits (Table 3). However, our data suggest that the LREEs are particularly depleted compared to the HREEs (Ce/Yb ~0.1), indicating


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Page 17 of 21 1079

a lack of the flat REE pattern (Ce/Yb >1) suggested pre-viously as evidence for interaction of peridotites with sub-stantial amounts of hydrothermal fluids (e.g., Boschi et al. 2006a). The trace element compositions indicate enrich-ments in Ba, K and Pb (Fig. 6) compared to DMM (after Salters and Stracke 2004). In contrast, Sr concentrations are very low compared to most mid-ocean ridge serpent-inites (Fig. 6d: e.g., from the Atlantis Massif, the 15°20′N FZ; Paulick et al. 2006; Delacour et al. 2008c). The low Sr concentrations support the observations discussed above that water–rock ratios were low and that seawater introduc-tion was limited.

Analyzed Pb concentrations are particularly high (Fig. 6c), with the unusual enrichment of Pb most likely reflecting either the abundance of sulfide phases (in which Pb, being a chalcophile element, is preferentially incorpo-rated) or resulting from high-temperature hydrothermal fluids (Staudigel 2003). Furthermore, Ba is also slightly enriched compared to DMM, especially in two samples (Fig. 6b). Ba is relatively immobile during serpentiniza-tion, but highly mobile during hydrothermal alteration, e.g., during massive sulfide mineralization (Niu 2004). Accord-ingly, high-temperature vent fluids such as at Rainbow, Broken Spur, TAG and Lucky Strike typically have strongly elevated Ba contents (von Damm 1990; James et al. 1995; Douville et al. 2002). Compared with other partly serpenti-nized peridotites, the Santa Elena samples have Ba contents similar to those from Logatchev and most samples from the 15°20′N fracture zone (Fig. 6b). Importantly, some serpen-tinites from the 15°20′N FZ preserve high alteration tem-peratures (>350 °C; Alt et al. 2007), which are generally associated with elevated Ba contents (Fig. 6c; Paulick et al. 2006; Kodolanyi et al. 2012). The relatively small amount of Cu-bearing phases in the Santa Elena peridotites implies that interaction with high-temperature fluids was signifi-cantly less extensive than at Logatchev and Rainbow, or in silica metasomatized serpentinites from the 15°20′N FZ or the Atlantis Massif. This possibly also explains the low REE concentrations. Additionally, substantial alteration of REE compositions during hydrothermal alteration requires high (>102) water–rock ratios (Bau 1991). Thus, the impact of the high-temperature fluid on the bulk rock trace ele-ment composition of the Santa Elena peridotites was prob-ably only minor, but the variably elevated Cu, Zn, Ba and Pb contents record evidence of this limited hydrothermal activity.

Summary and concluding remarks

The peridotites of the Santa Elena Ophiolite contain min-eral assemblages that are characteristic for serpentiniza-tion of ultramafic rocks at highly reducing conditions,

temperatures <250 °C and very low water–rock ratios. As discussed, two models may explain the formation of Cu-bearing sulfides and native copper: (1) the Cu-bearing phases formed through breakdown of primary sulfides due to highly reducing and sulfur-depleted serpentiniza-tion conditions at approximately 200–250 °C, or (2) they formed through introduction of a Cu-bearing, higher tem-perature fluid (350–400 °C) postdating serpentinization, similar to processes today occurring at the Logatchev and Rainbow hydrothermal fields. Model 1 does not well explain observations in thin sections, as more native copper is present than even in the most conservative mass balance calculations performed here. Unless there was significant chalcopyrite present as a primary sulfide (which all had to be replaced during desulfurization) or Cu was soluble dur-ing desulfurization, it is unlikely that the studied Cu-bear-ing phases simply formed by breakdown of the primary sulfides during serpentinization. Furthermore, simple des-ulfurization of primary sulfides cannot explain the higher temperature signatures preserved in awaruite–pentlandite pairs, unless serpentinization started at higher temperatures (350–400 °C) than predicted by the silicate mineralogy. However, the calculations performed in this study support mineralogical observations from other settings (e.g., the Zambales ophiolite) in which reducing and sulfur-depleted serpentinizing fluids resulted in breakdown of primary sulfides, forming bornite, chalcocite and native copper.

Consequently, we favor the second model and suggest that Cu-bearing sulfides and native copper in the Santa Elena peridotites were produced primarily through inter-action with a high-temperature fluid after serpentiniza-tion. Even though Cu contents are significantly lower than in typical black smoker type systems, they are within the range of partly serpentinized peridotites from Rainbow and Logatchev. The similarities to these systems, coupled with elevated Cu, Zn, Ba and Pb contents, suggest that the Santa Elena peridotites experienced similar processes to the base-ment of the Rainbow or Logatchev hydrothermal systems, but at a less extensive scale. This is consistent with the low degrees of serpentinization in most samples, the lack of pronounced Sr enrichments and the fact that low water–rock ratios (<1) resulted in minor modification of bulk rock trace element geochemistry (e.g., the LREE) by the fluids. Accordingly, we suggest that the primary minerals domi-nate the bulk rock geochemistry.

The Cu in these peridotites could have been sourced from layered gabbros and diabase dikes today exposed on the Santa Elena peninsula. No extensive Cu–Fe–Ni sulfide deposits (such as stockwork or massif sulfide deposits) are preserved on the Santa Elena peninsula, but the upper sec-tions of a possible hydrothermal system may have been removed during tectonic emplacement, erosion and/or weathering. Further evidence that the studied peridotites


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preserve the lower section of a hydrothermal system is also given by the unusually low water–rock ratios and strongly reducing conditions, reflecting strongly limited seawater introduction. Additionally, direct exposure of the peridotites to seawater would most likely have resulted in the formation of carbonate veins, as observed in seawater-exposed peridotites (Bach et al. 2011; Schwarzenbach et al. 2013b), but which are absent throughout the stud-ied samples. This suggests that the peridotites were not directly exposed to seawater, but were serpentinized at depth, presumably in a Cretaceous hydrothermal system along an oceanic spreading center during the opening of the Caribbean Sea that separated North and South America (Pindell et al. 2006). We therefore suggest that the struc-tures and mineral assemblages of the Santa Elena perido-tites provide a unique insight into the hydrothermal evolu-tion of peridotite-hosted hydrothermal systems, revealing the processes that occur within the deeper sections of the oceanic lithosphere.

Acknowledgments We thank J. Beard for motivating discussions and J. Snow for providing additional samples. S. Mazza, W. Whalen and H. Brooks helped with sample preparation and analytical work, L. Fedele and R. Tracy helped with EMP analyses. The authors acknowl-edge the valuable cooperation of the Area de Concervacion Guan-acaste, especially R. Blanco Segura (Research Program Coordinator) and M. M. Chavarría (Biodiversity especialities). Field assistance and participation by P. Madrigal, J. Calvo, M. Loocke and S. Wright was fundamental for field expeditions. Logistics and intellectual collabo-ration with P. Denyer (Central American School of Geology, UCR) was key for this project. We also thank O. Müntener, R. Frost and an anonymous reviewer for helpful comments that greatly improved the manuscript. This project was supported by the National Science Foun-dation award No. EAR-1019327 to Gazel. E.S. and M.C. gratefully acknowledge support from Virginia Tech Department of Geosciences.

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