de haller and fontbote-2009

0361-0128/09/3816/365-20 365

IntroductionTHE RAL-CONDESTABLE deposit is located 90 km south ofLima, Peru, approximately 5 km from the Pacific coast. Thedeposits estimated past production and current reserves exceed 32 million metric tons (Mt) at 1.7 percent Cu, 0.3 g/t

Au, and 6 g/t Ag. The ore typically consists of chalcopyrite,pyrite, pyrrhotite, and magnetite, found as disseminations,mantos, and veins in amphibolitized rocks belonging to aLower Cretaceous volcano-sedimentary sequence. Two maincontrasting genetic models have been proposed for this deposit. In the first, ore was interpreted to have been depositedsyngenetically through volcano-exhalative processes at or nearthe sea floor (Ripley and Ohmoto, 1977, 1979; Wauschkuhn,1979; Cardozo, 1983). The second view considers that the

The Ral-Condestable Iron Oxide Copper-Gold Deposit, Central Coast of Peru: Ore and Related Hydrothermal Alteration, Sulfur Isotopes,

and Thermodynamic Constraints*

ANTOINE DE HALLER,1 AND LLUS FONTBOT2

1University of Geneva, Mineralogy Department, rue des Marachers 13, CH-1205 Genve, Switzerland, and University of Bern, Institute of Geological Sciences, RWI Group, Baltzerstrasse 1-3, CH-3012 Bern, Switzerland

2University of Geneva, Mineralogy Department, rue des Marachers 13, CH-1205 Genve, Switzerland

AbstractThe iron oxide copper-gold (IOCG) Ral-Condestable deposit is located 90 km south of Lima, Peru, and ap-

proximately 5 km from the Pacific coast. Mineralization consists mainly of replacement mantos and dissemi-nations within permeable volcaniclastic and carbonate-rich rocks and structurally controlled veins surroundinga coeval and apparently causative intrusion of tonalitic composition emplaced in the core of a dacitic volcano.Potassic (biotite grading upward to sericite-chlorite) alteration and a poorly developed, almost sulfide-free,quartz stockwork closely border the tonalite, affecting the basaltic to dacitic Lower Cretaceous volcano-sedi-mentary host sequence. Ore is associated with a hydrated calc-silicate (mainly amphiboles) alteration that sur-rounds the biotite alteration. A hematite-chlorite (albite, epidote, calcite) alteration affects the periphery of thesystem. The main ore stage is characterized by two end-member mineral associations that were formed ac-cording to (1) an oxidized deposition sequence (hematite-magnetite-pyrite-chalcopyrite) occurring in and nearfeeder structures, and (2) a reduced deposition sequence (pyrrhotite-pyrite-chalcopyrite) found in volcaniclas-tic rocks and veins. Early specular hematite of the oxidized sequence is transformed to magnetite (mushke-tovite). The main ore-stage mineralization is cut by minor late-stage calcite-sulfide veins.

Main ore-stage sulfides have 34S values asymmetrically distributed from 1.0 to 26.3 per mil, with a medianat 6.6 per mil (n = 51). Similar values are observed for pyrrhotite, pyrite, and chalcopyrite. The 34S values de-pend on the stratigraphic position, with deep-seated vein samples normally distributed between 1.0 and 6.3 permil (avg about 3.5, n = 13) and shallower samples from 2.7 to 26.3 per mil (median around 7.5, n = 39).Sulfides found in late-stage calcite-sulfide veins show strongly negative 34S values ranging between 32.7 and22.9 per mil (n = 6), indicating a possible biogenic source.

Because no rock unit is known to occur in the internal parts of the deposit that could have oxidized fluids tothe point of hematite stability, the oxidized mineral sequence is best explained by magmatic brines followingthe SO2-H2S gas buffer at high temperature (>350C) and fluid/rock ratio. This is supported by the close tomagmatic 34S values of sulfides from the deep parts of feeder veins. Mass-balance calculation based on sulfurisotope data suggests that at the deposit scale, the bulk of the sulfides is dominated by magmatic sulfur, withsulfides of the oxidized minerals association having a larger component of magmatic sulfur than those of thereduced mineral association. The deposition sequence from hematite to chalcopyrite reflects the cooling of themagmatic fluid and redox and pH buffering by the basaltic-andesitic volcano-sedimentary host rocks. Thus, theoccurrence of magnetite pseudomorphous after early hematite (mushketovite) paragenetically followed byiron-bearing sulfides is interpreted to be direct field evidence for precipitation from oxidized magmatic brines.The same sequence has been described in many IOCG, skarn, and some porphyry copper deposits worldwide.34S values of sulfides ranging up to 26.3 per mil are found in what corresponded to a relatively shallow aquifer

filled with evolved reduced seawater. Heavy sulfur in H2S was produced through thermochemical reduction ofAptian seawater sulfate (34S = 14) in the recharge zone, which is interpreted to correspond to the hematite-chlorite (albite, epidote, calcite) alteration present at the upper flanks of the hydrothermal system, adjacent tothe causative intrusion. Hematitization (through oxidation) resulted from the high fO2 of seawater and from thereduction of its sulfate to H2S by the Fe2+ contained in the rock. In the core of the system, the seawater-derivedfluids reached near chemical equilibrium with their actinolitized host rock, at about 300 to 350C, in reduced,rock-dominated conditions. Mixing of these fluids with magmatic brines, already partially or totally reducedthrough reaction with wall rock at medium to low magmatic fluid/rock ratio can explain the large positive 34Sscatter observed in sulfides of the reduced mineral association, at stratigraphically shallow positions.

*Please refer to errata in this issue (p. 447) for corrections to deHaller etal. (2006).

Corresponding author: e-mail, [email protected]

2009 Society of Economic Geologists, Inc.Economic Geology, v. 104, pp. 365384

deposit formed by a single epigenetic hydrothermal process,which overprinted the whole volcano-sedimentary sequence.Atkin et al. (1985) and Vidal et al. (1990) proposed an am-phibolitic Cu-Fe skarn type for Ral-Condestable and otherdeposits located south on the Peruvian coast (Eliana, Mon-terrosas, Marcona, and Acari). More recently, Barton andJohnson (1996), de Haller (2000, 2006), de Haller et al. (2001,2002, 2006), Injoque (2002), Sillitoe (2003), and Williams etal. (2005) classified Ral-Condestable as iron oxide copper-gold type (IOCG), based on its mineral association, related al-teration, and style of mineralization (disseminations, mantos,and veins).

The genesis of IOCG deposits is still a matter of debate(e.g., Williams et al., 2005), and although it has been shownthat most Andean deposits are coeval with nearby plutonism(e.g., Marschik and Fontbot, 2001; Sillitoe, 2003; Marschikand Sllner, 2006), a direct linkage with a particular magmaticevent and its released magmatic fluids has proved difficult todemonstrate. This led to two main hypotheses for the originof the hydrothermal fluids: magmatic and nonmagmatic. Thefirst hypothesis states that the mineralizing fluids are mag-matic, and mixing with external fluids (seawater, evaporite-derived, metamorphic) might have happened but is not re-quired (Pollard, 2000, 2001; Sillitoe, 2003). In the secondcase, plutons only served as heat engines that drove convec-tion systems involving external brines which leached metalsfrom the country rocks during the heating path and precipi-tated ore in the upflow zone (Barton and Johnson, 1996,2000; Haynes, 2000; Hitzman, 2000). Hitzman (2000) consid-ered that the influx of external, nonmagmatic, saline, oxi-dized, and Cu-rich fluids is essential for the formation ofIOCG deposits, even in contexts similar to porphyry copper.A more recent third theory considers that, whereas metals areprobably derived from magmas, the input of sulfur from ex-ternal brines might be essential for sulfide precipitation (e.g.,Benavides et al., 2007).

In a previous paper, we showed, by geologic mapping andU-Pb dating of both intrusive and hydrothermal activity atRal-Condestable (de Haller et al., 2006), the existence of anintimate and well-exposed spatial connection and a remark-able temporal coincidence between tonalitic magmatism andcopper mineralization (coeval at 115 Ma within 2 error of

assessed by replicate analyses of laboratory standard materi-als (synthetic barium sulfate at +12.5 34S; natural pyrite at7.0 34S) was better than 0.2 per mil (2). The accuracyof the 34S analyses has been checked periodically by analysesof international reference standards: IAEA-S1 and S2 silversulfide (0.3 and +21.7, respectively), and NBS-123 zincsulfide (+17.3).

Hydrothermal AlterationAlteration types are summarized in Table 1 with their re-

spective mineralogy, style, position, and timing. The distribu-tion of the main alteration types is shown in the geologic sec-tion (Fig. 1).

The tonalite 1 stock is surrounded by a pervasive biotitealteration halo mostly developed in mafic rocks (Figs. 1, 3A-C). Biotite alteration is cut by a quartz stockwork and locallyby actinolite veinlets (

Magmatic amphiboles in quartz-diorite porphyry correspondto magnesiohornblende, whereas alteration amphiboles havecompositions centered on the actinolite-ferroactinolite-ferro-hornblende-magnesiohornblende intersection (Fig. 4B; mi-croprobe data in de Haller, 2006). Hydrothermal amphibolesare enriched in Fe and depleted in Ti and Mg compared tomagmatic hornblende. Like Cardozo (1983), we did not findanhydrous hydrothermal calcsilicates (garnet, pyroxene) inthe mined area. The only occurrence of garnet we found waspreviously described by Injoque (1985) and consists of a few-meters outcrop located about 2 km northwest of the Con-destable mine. It is unclear if this garnet is genetically relatedwith the iron-copper mineralization. Vidal et al. (1990) de-scribed such minerals, but based on the source work of In-joque (1985) it appears that garnet was not observed in themined area (except for a single occurrence of garnet in theCondestable mine), and pyroxene described in what corre-sponds to the subunit IIIB is possibly of magmatic origin(Cardozo, 1983). Ripley and Ohmoto (1977) did not describegarnet or pyroxene in the deposit, with the exception of localoccurrence of thin (15-cm) metamorphic halos of garnet andvesuvianite at dolerite dikes margin.

Actinolite alteration grades locally into magnetite alterationwhich affects predominantly the glass matrix, either of mafic

(unit III, Fig. 3F) or intermediate (quartz-diorite porphyry)volcanic and subvolcanic rocks. Mafic phenocrysts in unit IIIbasalt or basalt-andesite affected by magnetite alteration arecompletely actinolitized (Fig. 3F), whereas magmatic horn-blende in quartz-diorite porphyry remained intact. Lime-stones are replaced by specular hematite (Fig. 3K) proximalto feeder veins and by massive magnetite in more distal set-tings. Hematite is almost completely transformed into specu-lar hematite (mushketovite, Fig. 3K).

Actinolite and magnetite alterations grade upward intosericite-chlorite alteration (Fig. 3L), which affects the unit IVdacite-andesite dome, part of the underlying quartz-dioriteporphyry, and locally the top of subunit IIID (Fig. 1). The al-teration intensity is higher near feeder veins, where plagio-clase phenocrysts can be completely transformed to sericite.Magmatic hornblende phenocrysts and part of the volcanicglass are replaced by Fe chlorite.

Hematite-chlorite (with albite, epidote, calcite) alterationoccurs in the upper periphery of the hydrothermal system andsurrounds the sericite-chlorite alteration (Fig. 1). It affectsunit IV and the upper part of unit III, giving a characteristi-cally reddish to greenish color to the rocks. Magmatic plagio-clase can be totally albitized (e.g., sample AH-150, Fig. 4A),and chloritization and/or hematization of the ferromagnesianminerals is widespread. Locally, patches and veinlets of epi-dote and carbonate are present. Sericite is almost absent.

A mass-balance study of altered rocks of the Ral-Con-destable superunit magma series based on immobile ele-ments has been presented in de Haller (2006), together withthe details of the calculations. It showed that actinolite-al-tered rocks are enriched in Ca, Ti, and Mg and depleted in K,with albitized samples also showing Na gain and Ca loss.Sericite-chlorite alteration is associated with Fe, Mg, and Kgain and Ca loss, whereas Fe essentially remained immobilein the hematite-chlorite-albite alteration, which is character-ized by Mg and Na enrichment and Ca depletion.

The waning stages (postdating sulfide deposition) of the hy-drothermal system are characterized by the assemblageprehnite pumpellyite, which locally overprints actinolite al-teration in mafic rocks of unit III (Fig. 3J), and by sericite-chlorite, which affects (weakly) the core of the system, includ-ing the tonalite 1 stock. At the contact with dolerite dikes, athin (up to 1-m) alteration halo of epidote has been observed.

Ore Mineralogy, Distribution, and ParagenesisTwo superposed hydrothermal events are distinguished in

the deposit area: the main ore stage, which corresponds to thebulk of the iron oxide copper-gold mineralization, and late-stage calcite-sulfide veins of little economic importance.

The main ore-stage mineralization occurs on both sides ofthe 2 0.5 km tonalite 1 stock, in a distal position comparedto the quartz stockwork, mainly in unit III (Fig. 1; de Haller etal., 2006, figs. 3, 4). Copper ore is mainly associated with acti-nolite alteration and subordinately to sericite-chlorite. Themain ore stage is cut by a regional (tens of km along strike),northwest-trending dolerite dike swarm. Most of the ore con-sists of mantos (i.e., orebodies roughly conformable with bed-ding) and disseminations occurring in unit III around feederveins, as replacements and pore infillings of chemically reac-tive and/or porous beds consisting of carbonate rocks, tuff,

368 DE HALLER AND FONTBOT

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Thi

ckne

ssRock Types Mineralization

300

to 4

50 m

80 t

o 90

m ~

500

m27

0 to

360

m

limestone (bioclastic, upper beds are oolitic), shale, tuff,basalt to basalt-andesite

subarkose,tuff,lapillistone

basalt, volcanic breccia lenses,two limestone beds at bottom

basalt-andesite,volcanic breccia,tuff,sandstone

Sub

unit

A

B

C

D

Replacement and dissemination in tuffaceous beds

Dissemination and pore infills in volcanic brecciasDissemination in tuff

Disseminationin tuffReplacement of calcareous beds

D

Replacement of calcareous beds

pore infills in volcanic breccia

issemination and

UN

IT II

I

FIG. 2. Lithostratigraphic subdivisions and mineralization in unit IIIwithin the Ral-Condestable deposit (simplified from de Haller et al., 2006).

RAL-CONDESTABLE IOCG DEPOSIT, PERU 369

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TAB

LE

1. T

ypol

ogy

and

Rel

ativ

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hron

olog

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the

Alte

ratio

n of

the

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Min

eral

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Alte

ratio

n ty

peF

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c ro

cks

(Ra

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sup

erun

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afic

roc

ks (

unit

III)

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posi

tion

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ing

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anic

gla

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lcan

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lass

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ore:

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0 m

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o ar

ound

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arly

hig

h-te

mpe

ratu

re

mag

mat

ic h

bl

btm

afic

min

eral

s

btR

al-C

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stab

le s

uper

unit

aure

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arou

nd to

nalit

e to

nalit

e in

trus

ions

intr

usio

ns

Qua

rtz

stoc

kwor

kVe

ins:

qtz

a

ct

bt

ms

kfs

il

m,

Vein

s: q

tz

act

b

t k

fs

Stoc

kwor

k an

d ve

ins

Cor

e:

200

m a

ureo

le a

roun

d Sy

nore

: cut

s bi

otite

ttn

cp

tt

n

cpR

al-C

onde

stab

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uper

unit

alte

ratio

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lass

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c

p

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act

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the

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nd q

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(oxi

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ep

plag

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c

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ta, m

ain

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web

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; min

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abb

revi

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b =

albi

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and

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a am

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), bt

= b

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, ep

= ep

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, ilm

= il

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= m

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mag

netit

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= pr

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ump

= pu

mpe

llyite

, qtz

= q

uart

z, s

er =

ser

icite


0361-0128/98/000/000-00 $6.00 370

pyroclastic deposits, and volcanic breccias. Sedimentary fea-tures, such as thin bedding, are commonly preserved evenafter complete replacement (Fig. 3D, E, M). Mantos can ex-tend for up to about 200 m laterally with a 2- to 15-m thick-ness. Feeder veins are up to 4 m thick and cut the whole vol-cano-sedimentary sequence from unit II (at least) to unit IV,including the quartz-diorite porphyry sill-dike complex and,locally, tonalite 1. Whereas most of the veins in the Ral minefollow a northeast strike, the Condestable mine is dominatedby northwest-oriented veins. In both mines, the northwest-striking veins dip to the northeast, while the northeast-strik-ing veins are nearly vertical or steeply dip to the southeast(Fig. 5A). The displacements associated with veins are gener-ally minor (

3K). Later chalcopyrite filled the remaining porosity andpartly replaced the former pyrite.

A second, reduced mineral association was formed in thesequence: (1) pyrrhotite, (2) pyrite, and (3) chalcopyrite.Pyrrhotite is common in veins, volcanic breccias (subunit IIIC;Fig. 7D), and replacement mantos (subunit IIID). It com-monly contains tiny rounded inclusions of ilmenite and blebsof plumose quartz (full of microscopic fluid inclusions indi-cating recrystallization after amorphous silica or chalcedony).Pyrrhotite is locally associated with minor amounts of spha-lerite and molybdenite. In places, pyrrhotite is completely orpartly replaced by a pyrite-marcasite mixture (e.g., Ramdohr,1980; Vaughan and Craig, 1997; Fig. 7H-I). The later veiningand replacement of pyrrhotite by pyrite was followed by thedeposition of chalcopyrite as open-space fillings and replace-ments of the earlier pyrrhotite and pyrite (Fig. 7J).

Pyrite and chalcopyrite alone represent an intermediateredox mineral association. Magnetite and pyrrhotite are notfound together in the deposit (de Haller, 2006; M. Carpio andJ. Ziga, Ca. Minera Condestable S.A., pers. commun.,2005), except in rare occurrences where pyrrhotite occurs asminor (

that replaced bedded sediments of subunit IIID (Con-destable mine open pit).

In all three mineral associations, the deposition of pyriteand chalcopyrite was locally followed by sphalerite andgalena. Chalcopyrite contains traces of gold as inclusions(normally


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stockwork, itself containing biotite flakes, developed at the con-tact with the tonalite, suggests that magmatic fluids may havebeen released into the country rock through the hydraulicbreaching of a high-temperature, brittle-ductile carapace (cf.Fournier, 1999). It is considered unlikely that externally de-rived fluids, at or close to hydrostatic pressure, could havecrossed the brittle-ductile transition and entered the overpres-sured high-temperature region that surrounded the intrusion.Hypersaline fluid inclusions in the quartz stockwork containsignificant amounts of iron as chloride daughter crystals (deHaller, 2006) and traces of chalcopyrite are found intergrownwith the quartz, which suggests that the released magmatic flu-ids were metal-, sulfur-, and chlorine-rich, in agreement withpreliminary ICP-MS analysis on liquid-vapor fluid inclusionscontaining Cu, Zn, Pb, Au, and Ag from stockwork quartz (K.Kouzmanov, University of Geneva, pers. commun.).

In contrast to the potassic alterations, the shallow-lateralposition (Fig. 1) and mineralogy of the hematite-chlorite-al-bite alteration are consistent with a seawater recharge zone,in which the iron of the rocks was oxidized to hematite due tothe high fO2 of seawater combined with sulfate reduction toH2S (cf. Shanks et al., 1981; Reed, 1997). Albitization (Nainput) and chloritization (Mg input) are compatible with sea-water following a heating path toward the tonalite body in thedownflow zone (e.g., Giggenbach, 1984; Carten, 1986; Reed,1997). Present in deeper parts, the actinolite alteration is in-terpreted to correspond to the zone where reduced seawater-derived fluids were heated above 300C, in rock-dominatedconditions (cf. the lower temperature of actinolite stability inHenley and Ellis, 1983; Reed, 1997).

An input of meteoric water in the upper parts (eroded) ofthe system cannot be ruled out.

Sulfur isotopes

The 34S value of pyrrhotite reflects the isotopic composi-tion of the fluid because at the redox and pH conditions ofpyrrhotite precipitation (reduced and slightly acid), all thesulfur is in the form of H2S (Ohmoto, 1972). The 34S valueof pyrrhotite range from 3.7 to 21.1 per mil (n = 13, this studyand Ripley and Ohmoto, 1977; Table 3) with major variationseven at the sample scale, indicating that it precipitated fromfluids with widely varying 34SS (cf. Ohmoto and Rye, 1979;Ohmoto and Goldhaber, 1997). Similarly as during pyrrhotiteprecipitation, reducing (chalcopyrite and pyrite are com-monly associated with pyrrhotite or magnetite but not with

hematite) and slightly acid conditions apparently prevailedduring pyrite and chalcopyrite deposition and their 34S val-ues (1.026.3) most probably also nearly reflect the iso-topic composition of the fluid (34SS).

The large range of 34S values of sulfides points to the mix-ing of two or more fluids with different 34SS. Three possiblesulfur reservoirs can be considered in the context of the Ral-Condestable deposit, namely magmatic, seawater, and evap-orite deposits.

The isotopic composition of magmatic sulfur at Ral-Con-destable is assumed to fall in the 0 5 per mil 34S range, val-ues encountered in most skarn and porphyry copper depositsassociated with I-type granitoids in the Andes (Ohmoto andGoldhaber, 1997).

The isotopic composition of seawater sulfate was 14 1 permil at the Aptian-Albian boundary (Claypool et al., 1980).This value is much lower than the highest values of main ore-stage sulfides (e.g., up to 26.3, sample AH-16 pyrite; Fig.7H, Table 2). This situation is not common in ore depositsformed from seawater-derived sulfur (Huston, 1999; Shanks,2001), but thermochemical reduction of seawater sulfate by


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FIG. 7. Mineralization. A. Sample AH-56 of quartz stockwork in quartz-diorite porphyry. B. Sample AH-55 of quartzstockwork in biotitized basalt of subunit IIID. C. Sample AH-10 of the Argentina vein, Ral mine; massive milky quartz cutby chalcopyrite-sphalerite-chlorite veinlets in amphibolitized-chloritized basaltic host rock. D. Ral mine: concordant acti-nolitized basaltic volcanic breccia (subunit IIIC) filled by pyrrhotite with minor pyrite and chalcopyrite (photo by K. Kouz-manov). E. Sample JZ-16 of the Pampa vein ore; allanite crystals surrounded by pleochroic halos in hydrothermal actinolitein association with molybdenite and chalcopyrite (opaque). F. Sample AH-12 from the Ral mine; actinolitized subunit IIIAbasalt cut by veins filled with mushketovite, pyrite, and chalcopyrite. G. Sample AH-80 from the Ral mine; replacement oflimestone (subunit IIID) by mushketovite (magnetite after specular hematite), pyrite, minor chalcopyrite, and chlorite-acti-nolite. H. Sample AH-16 of massive pyrite ore from vein 53 (air-etched, reflected light view); py1 shows birds eye texture(Ramdohr, 1980), indicating that it probably formed from pyrrhotite sulfidization. Py1 is cut by a later py2 vein, followed bychalcopyrite deposition in porosity and cracks (see isotope results in Table 2). I. Sample AH-38 of massive pyrrhotite fromthe Chilena vein with pyrite-marcasite birds eye texture as in (H) (Ramdohr, 1980, see sulfur isotope results in Table 2). J.Sample AH-17 of massive pyrite ore from vein 53; cracks in pyrite are filled by late chalcopyrite. Abbreviations: act = acti-nolite, all = allanite, cp = chalcopyrite, i.p. = intermediate product (pyrite and marcasite from pyrrhotite sulfidization), mo= molybdenite, mt = magnetite, po = pyrrhotite, py = pyrite, qtz = quartz.

cp-py-(mt mk) py-mt

mt

mt

pycp

mk

Oxidized mineral association

pycp

pycp-py

Intermediate redox mineral association

cp-py py-po po

po (i.e.)cp py

Reduced mineral association

PROXIMAL DISTAL

FEE

DE

R V

EIN

host rock

host rock

host rock

FIG. 8. Main ore-stage mineralogical zoning. The oxidized, intermedi-ate redox, and reduced ore mineral associations are shown. These are notstable assemblages (in the sense of Einaudi et al., 2003) as these minerals didnot precipitate simultaneously (see Fig. 6). Abbreviations: cp = chalcopyrite,mk = mushketovite (magnetite replacing hematite), mt = magnetite, po =pyrrhotite, py = pyrite.


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TABLE 2. Results of Sulfur Isotope Analyses of Sulfides of the Ral-Condestable Deposit

Mine UTM UTM Redox4 34SSample Mine level 1 East North Sample description 2 Unit3 min. assoc. Mineral2 ()5,6

Main ore stageAH-9 Ral -30 327641 8595307 Vein infill: massive scp with cp + py + ser IIIC i cp 6.9

+ Kfeld + chl ( + sl + Au) i py 6.6AH-11 Ral -120 327770 8595154 Vein infill: massive scap with cp + sl + gn IIIC i cp 3.9

+ Kfeld + ttn veinlets i py 4.9AH-12 Ral -140 328270 8594798 Altered lava (act + chl + ap) with mt IIIA o cp 5.8

(after hm) + cp + py veins o py 6.3AH-13 Ral 0 327733 8595209 Altered lava (chl + Kfeld) with disseminated IIIC i cp 7.9

cp + py + sl, cut by late cal veinAH-14 Ral -30 328035 8595011 Vein infill: cp + sl and late cal IIIC i cp 10.2

i sl 10.2AH-16 Ral 0 327586 8595383 Vein infill: massive botryoidal py1 IIID r py1 26.3

(+ mt + cp); py2=late veinlets r py1 17.8r py2 6.3r py2 6.8

AH-22 Ral -30 328041 8595011 Altered volcanic breccia (act + ilm + ser + IIIC r cp 20.1Kfeld) with cp + py ( + po) r py 20.2

AH-37 Ral -30 328035 8595011 Altered lava (preh + pum + chl + act + ttn) IIIC i cp 4.4with cp + cal diss. and vein

AH-38 Ral -30 327777 8595114 Vein infill: massive po + qtz + i.p. IIIC r cp 12.0r po 17.4r po 21.1r po 15.6r i.p. 6.6r i.p. 6.4r i.p. 24.4

AH-80 Ral 0 327463 8594872 Metasomatized limestone: trem + chl + cal IIID o cp 9.0+ mt (after hem) + py ( + cp) o py 11.2

AH-81 Ral -65 327997 8594986 Metasomatized limestone: act + mt + py (+ po + cp) IIID r py 14.4AH-82 Ral -65 327997 8594986 Metasomatized limestone: act + chl + py IIID r cp 6.0

+ cp ( + po + mc) r py 13.9AH-83 Ral -65 327997 8594986 Metasomatized limestone: chl + py (+ po + mt) IIID r py 5.0AH-84 Ral -65 327997 8594986 Metasomatized limestone: py + chl IIID o py 6.5

(+ qtz + mt + cp + amph)AH-85 Ral -140 328270 8594798 Altered lava (act + chl + ap) with cp + mt IIIA o cp 5.6

(+ py) veins o py 3.5AH-87 Ral -140 327752 8595169 Vein infill: massive po (+ act + qtz + cp + py + ilm) IIIC r po 7.3AH-88 Ral 60 327785 8595122 Altered lava (act) with cp + sl + gn + py vein IIIC i cp 8.7

i sl 9.2AH-89 Ral -140 327752 8595169 Vein infill: scp + act + chl + po + py (+ cp + ilm) IIIC r po 7.6

r py 7.3AH-92 Ral 0 327463 8594874 Metasomatized limestone: act + chl + cp + py IIID i cp 5.5AH-93 Ral -65 327276 8595008 Altered rock (act + chl) with py (+ cp + mt) IIID o py 2.7

as dissemination and vein o py 3.8AH-96 Ral -30 327641 8595307 Vein infill: scp + cp + mo IIIC i mo 7.0

i mo 6.9i cp 6.6

JZ-16 Ral -140 327997 8594986 Altered lava (act + chl + ap + ilm) with mo IIIB r mo 2.8+ cp (+ po + py) vein r cp 3.2

JZ-18 Ral -140 328210 8594940 Altered lava (act) with cp + py ( + sl) vein IIIA i cp 4.0i py 3.8

AH-33 Condestable -290 328100 8596350 Vein breccia infill: scp, ab, cp, po, py, cal, IIIA r po 4.0Kfeld, amph, chl, qtz, rock clasts r cp 1.0

AH-34 Condestable -230 327800 8596100 Vein infill; py ( + sl + cp + cal) IIIA i py 3.7AH-35 Condestable -230 328150 8596300 Altered lava (Kfeld + chl) with disseminated IIIA i py 1.4

cp + py + mo i cp 1.2

Late-stage calcite-sulfide veinsAH-6 Ral 60 327777 8595134 Late vein infill: banded cp + sl + cal IIIC cp -32.7

(+ cv + tn + py) sl -29.7AH-13 Ral 0 327733 8595209 Late veinlet cutting IOCG ore: sl + gn + py IIIC gn -28.6

+ mc + cal ( + cp) py -22.9AH-18 Ral 0 327586 8595383 Late vein infill: sl + py + mc + cal ( + cp) IIID sl -31.1

py -29.9

1 The 0-m levels at Ral and Condestable mines correspond to altitudes of 128 and 350 m, respectively2 Abbreviations: act = actinolite, amph = amphibole, Au = gold, cal = calcite, chl = chlorite, cp = chalcopyrite, cv = covellite, gn = galena, hm =

hematite, ilm=ilmenite, i.p. = intermediate product (mixture of marcasite and pyrite, Ramdohr, 1980), Kfeld = K-feldspar, mc = marcasite, mo = molyb-denite, mt = magnetite, po = pyrrhotite, preh = prehnite, pum = pumpellyite, py = pyrite, qtz = quartz, scp = scapolite, ser = sericite, sl = sphalerite, tn =tennantite, trem = tremolite

3 Refers to the lithostratigraphic units defined in Figure 34 Redox mineral associations: o = oxidized (mt present, no po), i = intermediate (without mt or po), r = reduced (po present)5 Values are given in per mil relative to CDT (Caon Diablo Troilite)6 Values in normal text correspond to the 2000 batch, and in italics to the 2004 batch (see text); analyses performed at the Stable Isotope Laboratory of

the University of Lausanne, Switzerland

the Fe2+ contained in the host rock in a system partially opento H2S (Shanks and Seyfried, 1987; Ohmoto, 1996) could ex-plain the few cases with sulfide 34S values higher than that ofcontemporaneous seawater sulfate (Fig. 12; Shanks et al.,1981; Zheng and Hoefs, 1993; Ohmoto and Goldhaber,1997). This hypothesis is supported by the 34S spread tohighly positive values (up to 26.3: Fig. 10) in the volcano-sedimentary subunits IIIC and IIID, characterized by muchhigher permeability than the deeper subunits IIIA and IIIB(Figs. 2, 7D; Table 4). Subunits IIIC and IIID apparently

acted as aquifers that allowed the lateral influx of seawater-derived fluids. In contrast, 34S values close to magmatic arerecorded in sulfides of the underlying subunits IIIA and IIIB.

Although evidence of evaporitic beds was not found in themapped area, regionally, Palacios et al. (1992) and Salazar andLanda (1993) described minor local gypsum occurrences inthe Morro Solar Group (equivalent to unit II) and in the Pam-plona Formation (equivalent to part of unit III; Fig. 1). Theoldest known rocks of the central Peruvian coast are of LateJurassic age (unit I, Fig. 1; or Puente Piedra Group in Rivera


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-15 0 +15 +3034S ()

gn (n = 1)sl (n = 4)cp (n = 19)py (n = 25)po (n = 6)mo (n = 2)

-30

Freq

uenc

y

0

15

5

10

20

Late stage calcite-sulfide veins Main ore stage sulfides

FIG. 9. Histogram of the sulfur isotope results. Mineral abbreviations: cp = chalcopyrite, gn = galena, mo = molybdenite,po = pyrrhotite, py = pyrite, sl = sphalerite. Data are in Table 2.

TABLE 3. Mean and Ranges of Sulfur Isotope Data for Sulfides from Ral-Condestable ()

Robust median Error34S abs. Confidence1 n Minimum Maximum

Main ore-stage sulfides (this work)py 6.5 +4.7/1.6 96.5 23 1.4 26.3cp 5.9 +2.8/1.9 96.9 18 1.0 20.1po 11.6 +9.5/7.6 96.9 6 4.0 21.1mo 2 2.8 7.0sl 2 9.2 10.2All sulfides 6.6 +1/0.8 95 51 1.0 26.3

Main ore-stage sulfides (data from this work, combined with those of Ripley and Ohmoto, 1977)2py 6.30 +1.5/0.50 95 53 1.4 26.3cp 5.35 +0.25/0.55 95 176 9.3 23.3po 7.40 +8.2/3.4 97.8 13 3.7 21.1sl 11.50 7.7 95 3 9.2 15.0gn 2 11.0 11.2All sulfides 5.70 +0.40/0.30 95 249 9.3 26.3

Late-stage calcite-sulfide veins (this work)py 2 29.9 22.9cp 32.7 1gn 28.6 1sl 2 31.1 29.7All sulfides 29.8 +6.9/2.9 96.9 6 32.7 22.9

Late-stage calcite-sulfide veins (data from this work, combined with those of Ripley and Ohmoto, 1977)2py 2 29.9 22.9cp 2 32.7 18.4gn 26.4 +16/2.2 93.8 5 28.6 10.0sl 26.1 +8.1/5.1 96.9 6 31.1 17.9All sulfides 26.4 +8.0/3.3 96.5 15 32.7 10.0

1 In percent2 Only the samples classified in the H category of mineral occurrence (sl + gn + cp veins) in Ripley and Ohmoto (1977), and two galena samples classi-

fied in the E stringer ore category with 34S values

et al., 1975; Palacios et al., 1992; Salazar and Landa, 1993),thus limiting the possible 34S of any evaporitic source to 14 to17 1 per mil (Claypool et al., 1980). The value of +23 per

mil used in the genetic model of Ripley and Ohmoto (1977)for seawater or evaporite sulfate is out of this range and farabove the Aptian seawater sulfate value. In the geologicallyunlikely case that evaporitic sulfur was involved, a similar rea-soning as for seawater sulfate reduction can be applied to ex-plain the heavy sulfide 34S values at Ral-Condestable.

We propose that a deeply sourced fluid dominated by mag-matic sulfur (avg 34Sfluid between 34: values for veins in


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TABLE 4. Sulfur Isotope of Main Ore-Stage

Subunit IIIA1 Subunit IIIB1

n Min. Max. Average4 Median 5 n Min. Max. Average4 Median5

Main ore-stage sulfides2Veins 11 1.0 6.3 3.7 1.2 3.8 +1.8/2.4 (93.5) 2 2.8 3.2 3.0 2.5Mantos and

disseminationsAll mineralized

structures 11 1.0 6.3 3.7 1.2 3.8 +1.8/2.4 (93.5) 2 2.8 3.2 3.0 2.5

Chalcopyrite3Veins 5 1 5.8 3.5 2.9 4.0 +1.8/3 (93.8) 7 0.8 8.1 4.1 2.2 3.4 +4.7/2.6 (98.4)Mantos and

disseminations 13 9.3 2.7 2.1 2.4 0.8 +2.5/5.2 (97.8) 58 0 8.3 4.11 0.50 4.0 +0.6/0.5 (95)

Pyrite3Veins 5 1.4 6.3 3.7 2.2 3.7 +2.6/2.3 (93.8)Mantos and

disseminations 5 4.1 8.3 5.2 2.2 4.6 +3.7/0.5 (93.8)

Pyrrhotite3Veins 1 4.0 4.0Mantos and

disseminations

1 Lithostratigraphic subunits used in this paper can be correlated to the units defined by Ripley and Ohmoto (1977), with subunit IIIA being equivalent to their unit I, subunit IIIB to their unit II, subunit IIIC to their units III and IV, and subunit IIID to their unit V

2 Data from this work (Table 2)3 Our results combined with those of Ripley and Ohmoto (1977)4 Absolute error at 95% confidence5 Error confidence percentage in parenthesis

-15 0 +15 +3034S ()

0

5IIIA

(0/11)

IIIB (0/2)0

15

5

10

IIIC (3/22)

0

5IIID (9/5)

allsulfides veins

replacements and disseminations

(1/4) number of samples of each ore type

IIIA lithostratigraphic subunit (see text)

FIG. 10. Histograms of the sulfur isotope results of the main ore-stage sul-fides by lithostratigraphic subunit and ore type. Data are in Table 2.

0.0 5.0 10.0 15.0 20.0 25.0 30.0

34Ssulfides

reduced

oxidized

intermediate

FIG. 11. Distribution of the 34S values of sulfides of the reduced(pyrrhotite-pyrite-chalcopyrite), intermediate (pyrite-chalcopyrite), and oxi-dized (hematite-magnetite-pyrite-chalcopyrite) ore mineral associations.Data are in Table 2.

subunits IIIA and IIIB) circulated upward in the feederveins. At shallower levels, it mixed with a seawater-derivedfluid in the permeable subunits IIIC and IIID. Mixing ofvarying amounts of these two fluids can explain the large iso-topic variations in sulfides from subunits IIIC and IIID, bothat the deposit and hand sample scale. Heavy 34S values,which indicate a larger proportion of seawater-derived sulfur,are mostly found in the reduced mineral association (Fig. 11),in subunits IIIC and IIID.

Assuming that the 34S of magmatic sulfur was 0 per miland that of seawater-derived H2S was 26 per mil (max valueon pyrite, Tables 2, 3), the median 34S value of the main ore-stage sulfides at 5.7 per mil (this work and Ripley andOhmoto, 1977; Table 3) would correspond to a mixture of 78percent magmatic- and 22 percent seawater-derived sulfurfor the bulk of the sulfides. If the Aptian seawater 34S valueof 14 per mil (Claypool et al., 1980) is taken as a minimum forseawater-derived H2S, this would correspond to a mixture of59 percent magmatic- and 41 percent seawater-derived sulfur. This suggests that the Ral-Condestable sulfides weredominated by magmatic sulfur and that external sulfur inputmight have been limited and even nonexistent in the deepparts of the system where sulfides with 34S values close tomagmatic are found (Fig. 10, Table 4). It follows that mixingwith the seawater-derived fluid (i.e., input of external sulfur)was apparently not required for sulfide deposition. A modelof the Ral-Condestable hydrothermal system showing thepossible sources of fluids is given in Figure 13.

Sulfides from late-stage calcite-sulfide veins have stronglynegative 34S values (32.7 to 22.9: Fig. 9, Table 3), sug-gesting recycling of biogenic sulfur leached from sediments


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Sulfides by Lithostratigraphic Units ()

Subunit IIIC 1 Subunit IIID 1

n Min. Max. Average4 Median5 n Min. Max. Average4 Median5

22 3.9 24.4 9.6 2.4 7.3 +2.9/0.7 (94.8) 5 2.7 26.3 12 12 6.8 +20/4.1 (93.8)

3 7.9 20.2 16 18 9 3.8 14.4 8 3 6.5 +7.4/1.5 (96.1)

25 3.9 24.4 10.4 2.4 7.6 +2.6/0.7 (95.7) 14 2.7 26.3 9.7 3.8 6.7 +7.2/1.2 (94.3)

15 3.9 20.3 8.6 2.5 8.2 +3.7/3.5 (96.5)

29 7.9 23.3 12.4 1.7 10.5 +3.9/1.1 (95) 49 0.6 12.7 5.54 0.55 5.4 +0.2/0.3 (95)

10 4.9 24.4 10.3 4.1 8.6 +4.8/2.2 (97.9) 4 6.3 26.3 14 15

3 14.5 23.1 19 11 26 2.7 14.5 7.5 1.4 6.2 +2.2/0.8 (95)

5 7.3 21.1 13.8 7.6 15.6 +5.5/8.3 (93.8)

7 3.7 8.2 6.2 1.7 6.6 +1.6/2.9 (98.4)

00.250.50.751

fraction of aqueous sulfate

34

S

(R) SO

200C

4

(R) SO 400

C4

(R) HS 400C2

Aptian Seawater sulfate

magmatic sulfur

S 34 range of main ore stage sulfides (this work)

-20

-10

0

10

20

30

40

50

60

70

80

90Seawater sulfate reduction

SO 200C

4

H S 400C

2

H S 200C

2

rang

e of

pos

sib

le

S

of s

eaw

ater

-der

ived

HS

34

2

(R)

HS

200

C2

S range 34

(subunits IIIA and IIIB)of deep sulfides

SO 400C

4

S range 34

(subunits IIIC and IIID)of shallow sulfides

FIG. 12. Fractionation of sulfur isotopes during seawater sulfate reduc-tion. Curves for closed system (equilibrium) and Rayleigh (R) fractionationare shown for 200 and 400C. Calculation was done with the sulfide-sulfatefractionation parameters given in Ohmoto and Goldhaber (1997). The 34S is14 per mil for Aptian seawater sulfate (Claypool et al., 1980) and 0 per milfor magmatic sulfur. The right side of the diagram corresponds to reducedconditions, where the 34S of the seawater-derived H2S stays between the ini-tial seawater sulfate value (total reduction at equilibrium) and any highervalue reached through Rayleigh fractionation. In Ral-Condestable, mag-matic sulfur is dominant in deep sulfides (subunits IIIA and IIIB), whereasinput of seawater-derived sulfur is significant in shallower sulfides (subunitsIIIC and IIID; see Fig. 10).

or bacterial reduction of seawater sulfate at temperaturesbelow 100C (e.g., Ohmoto and Goldhaber, 1997; de Halleret al., 2002). The fact that some late-stage calcite-sulfide veinscut post-main ore-stage dolerite dikes could suggest that twoseparate hydrothermal events with no genetic relationshipwere superposed, as previously proposed by Ripley andOhmoto (1977). However, similar late mineral associationshave been documented in other Andean IOCG deposits (e.g.,Candelaria-Punta del Cobre District, Chile), suggesting thata genetic connection between the two hydrothermal stages atRal-Condestable cannot be ruled out.

Probable precipitation mechanisms and genetic significanceof the oxidized and reduced mineral sequences

Actinolite alteration suggests a minimum temperature ofabout 300C (e.g., Henley and Ellis, 1983) for ore deposition,at a near-neutral pH (~5.6 at 300C; see Reed, 1997, fig. 7.6).

Fournier (1999) estimated the lower limit of the brittle-duc-tile transition in magmatic hydrothermal systems to be atabout 400C, which we interpret as a likely upper tempera-ture of ore deposition in Ral-Condestable. A temperature of350C is considered as an acceptable estimate for the seawa-ter-derived fluid in the core of the hydrothermal system. Pre-cipitation of pyrrhotite in a system which redox state is rockbuffered implies a minimum temperatures of about 230 to330C (mt + py/po line in Fig.14, considering RH up to 2.3for pure basalt; Einaudi et al., 2003), a 300 to 400C tem-perature range is considered as realistic for the deposition ofthe main ore-stage minerals (Fig. 6).

At a constant sulfur activity (i.e., broadly constant amountof sulfur in the system) and low sulfur/iron ratio, and con-sidering that pH is buffered by the wall rock, oxides are thestable iron minerals at high temperature, while sulfides ap-pear at lower temperatures (Fig. 15), and the precipitation


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?

hm-chlser-chl

qtz-bt

act+mt

act+mt

ser-chl

hm-chl

?

?

? ?

?

? ?

hm-chl

hm-chl

? ?

500

m

500 m

-1 km

-2 km

-3 km

-4 km

-5 km

0 km

+1 km

seawater level?seawater level?

Subunit IIIC

Subunit IIIC

? ?

?

Interpreted Cu anomaly

ser-chl

hm-chl

Alterationboundary

Vein

Inferred vein

Fault

Tonalite 1

Qtz-diorite porphyry

Tonalite 2

IV: volcanic

III: volcano-sed.

II: detrital

I: volcano-sed.

V: volcanicLi

tho

stat

igra

phi

c un

its

Intr

usiv

e ro

cks

seawatermagmatic fluidmeteoric fluid

Aptian seaAptian sea

FIG. 13. Fluid flux model of the Ral-Condestable hydrothermal system. West-southwesteast-northeastinterpretedcross section equivalent to the one shown in Figure 1 but in which geology has been rotated to its position at the time of min-eralization. The position of the seawater level is tentative. Evidence for meteoric fluids having been involved is lacking butcannot be excluded in the upper part of the system, now eroded. The highly permeable volcanoclastic subunit IIIC is shownwith a specific pattern. The deep part of the section is hypothetical.

of pyrite after pyrrhotite is compatible with a temperaturedrop under reduced conditions. Therefore, buffering atnear-neutral pH of a magmatic fluid by the host rocks dur-ing temperature decrease is considered as the more likelyprecipitation mechanisms for both oxides and sulfides inRal-Condestable.

At an Andean scale, Ral-Condestable in Peru and Cande-laria-Punta del Cobre in Chile are the only IOCG depositswhere both hematite and pyrrhotite have been described

(Sillitoe, 2003). In contrast, the oxidized mineral association isfound in most of the deposits, the replacement of hematite bymagnetite (mushketovite) recording a drop in the redox stateof the fluid (Figs. 3K, 6). Interestingly, the same sequence isalso widely documented in skarns (e.g., Boni et al., 1990;Dnkel, 2002; Ciobanu and Cook, 2004; Prendergast et al.,2005) and in some porphyry copper deposits, all being noto-riously of magmatic hydrothermal origin (e.g., Frikken et al.,2005). Magnetite (mushketovite) pseudomorphous after


0361-0128/98/000/000-00 $6.00 381

R =

log(

fH/f

HO

)H

22

100 200 300 400 500 600 800 1000-8

-7

-6

-5

-4

-3

-2

-1

3.0 2.8 2.6 2.4 2.2 2.0 1.8 1.6 1.4 1.2 1.0 0.8

1000/T [K]

T [C]

mt + S po

hm + S

py

mt + S

py

mt + py

pomt + qtz

fayalite

NiO Ni

plutonic and volcanic rocksin arc setting

brit

tled

uctil

e

hm (3)mt

hm (2)mt

hm (1)mt

(FeO )

(FeO)1.5

S-gas buffer (1 bar)

SOH

S2(g)2

(g)

SOH

S2(g)2

(g)

seawater-derived fluid

actin

olite

sta

ble

magmatic flu id

HSOH S

4

2 (aq)

-

(pH=6)

SOH S

4

2 (aq)

--

(pH=3)

oxi

diz

ed m

iner

alas

soci

atio

nre

duce

d m

iner

al

asso

ciat

ion

S-gas buffer (brine, 1000 bars)

Giggenbach rock buffer

FIG. 14. RH-1000/T diagram showing the phase boundaries in the system Fe-S-H2O and the position of the main redoxbuffers (modified from Einaudi et al., 2003). RH = log (fH2/fH2O). Mineral reactions were calculated at 500 bars. The positionof the hematite-magnetite phase boundary (1) proposed by Einaudi et al. (2003) was calculated from thermodynamic datapublished by Myers and Eugster (1983) for the 670 to 1,300C temperature range. However, extrapolation to temperatureslower than 670C is not recommended, and Myers and Eugster (1983) proposed other parameters for the 25 to 670C tem-perature interval, corresponding to the curve (2), where the stability field of hematite at hydrothermal temperatures is greatlyincreased compared to (1). By comparison, the position (3) of this phase boundary obtained using the program GeochemistsWorkbench and the included thermo.com.v8.r+.dat database (valid from 0 to 300C at 1 atm below 100C and along thewater vapor pressure at higher temperatures; thermodynamic properties for hematite and magnetite from Helgeson et al.,1978) lies at about 0.8 RH unit below the Myers and Eugster (1983) curve, with a similar slope. The difference between thepositions of curves (2) and (3) has little impact on the conclusions. The field of arc plutonic (upper left portion) and volcanicrocks (lower right portion) is shown in gray. The rock buffer and the position of the pressure-dependent S gas buffer arefrom Giggenbach (1987). A minimum temperature for the brittle-ductile limit has been indicated at 400C. The large un-filled arrow represents the general path followed by the magmatic fluid, which first follows the S gas buffer at high temper-ature and water/rock ratio, and then reacts at lower temperatures with the reduced wall rock so that the RH drops to the rockbuffer line. The gray arrow shows the approximated redox-T heating path followed by seawater, which progressively reactsand equilibrates with the host sequence. Because the seawater-derived fluid is reduced and at equilibrium with the rockbuffer at high temperatures (>200C), mixing with the magmatic fluid would have nearly the same effect as the reactionwith wall rock. Depending on the magmatic fluid/rock ratio and importance of fluid mixing, the mineralizing magmatic fluidfollows cooling trends between the upper oxidized path that reaches the hematite field (high magmatic fuid/rock ratio, lowmixing) and the lower reduced path that enters into the pyrrhotite field (low magmatic fluid/rock ratio, high mixing). Thesetwo end-member paths correspond respectively to the oxidized and reduced ore mineral associations described in Figure 6.Mineral abbreviations: hm = hematite, mt = magnetite, po = pyrrhotite, py = pyrite, and qtz = quartz.

hematite or hematite (martite) pseudomorphous after mag-netite is controlled by the redox reaction:

3Fe2O3(hm) + H2(aq) = 2Fe3O4(mt) + H2O, (1)

which is equivalent to

2Fe3O4(mt) + 12O2(aq) = 3Fe2O3(hm). (2)

As already pointed by Giggenbach (1987), the second reac-tion is conceptually correct but has no meaning in terms ofreal reactions because fluids near the hematite-magnetite sta-bility boundary at temperatures less than 500C contain vir-tually no O2. The hematite-magnetite pair therefore relates tothe H2 content of the fluid, as in reaction (1).

Ohmoto (2003) and Otake et al. (2007) proposed a nonredoxmechanism for the transformation of magnetite and hematitein hydrothermal systems based on the following reaction:

Fe2O3(hm) + Fe2+ + H2O = Fe3O4(mt) + 2H+. (3)

Although this reaction is valid at the redox boundary be-tween the fields of hematite and magnetite (and is therefore aredox buffer), its extension to higher or lower redox conditions(i.e., in the fields of hematite or magnetite stability) wouldimply metastable transformations not documented in naturalmineral assemblages. Additionally, Mcke and Cabral (2005)pointed that nonredox magnetite-hematite transformation im-plies important volume changes (>30%), incompatible withthe preservation of the pseudomorphic fabric of mushketovite.

In contrast, the redox transformation of hematite and magnetite(reaction 1) implies volume changes of less than 2 percent.

In the case of Ral-Condestable, no lithostratigraphic unitsare present in the internal parts of the deposit that could haveoxidized an early hydrothermal fluid to the point of hematitestability. On the contrary, a large part of units I and III (Fig. 1)consists of basalt to basalt-andesitic rocks, which lie near theFe2+/Fe3+ rock buffer line (Fig. 14; Giggenbach, 1987, 1992),and possibly down to about 0.5 RH unit below this line if purebasalts are considered (Einaudi et al., 2003). In the core of thehydrothermal system, a seawater-derived fluid (Figs. 13, 14)should have been reduced (i.e., on the rock buffer line), be-cause it must have been at redox equilibrium with its aquifer(mostly subunits IIIC and IIID; Fig. 2; e.g., Shanks et al., 1981;Ohmoto, 1986) at the temperatures (>300C) and lowwater/rock ratio suggested by the widespread occurrence ofactinolite alteration (Henley and Ellis, 1983; Reed, 1997).

Considering the deep-sourced fluid originated through thedegassing of a deep-seated intrusion at high temperature, itsredox state would have been controlled internally by the mag-matic SO2/H2S gas buffer (Fig. 14, Giggenbach, 1987, 1992,1997; Einaudi et al., 2003). This gas buffer is effective if fluidsare channeled upward through open veins at high flux andfluid/rock ratio, conditions that minimize the interaction withthe host rock. A temperature decrease following the SO2/H2Sbuffer line increases the relative oxidation state of the fluid, andthe hematite stability field is reached at temperatures between500 and 250C (Fig. 14; Helgeson et al., 1978; Myers and Eu-gster, 1983; Giggenbach, 1997; Einaudi et al., 2003). This ex-plains the presence of early hematite in the oxidized mineralassociation. Lateral and temporal zonation from early proximalhematite to distal and/or late magnetite together with the wide-spread transformation of hematite into magnetite are consis-tent with the magmatic fluid being progressively reducedthrough reaction with the wall rock. Acid generation throughthe dissociation of SO2 into HSO and H2S (and therefore wall-rock reactivity) increases at temperatures under about 350C,thus explaining the redox drop at this temperature (Fig. 14).

Whereas the oxidized sequence is interpreted to be indica-tive of magmatic-dominated fluids circulating at high fluid/rockratio, the reduced mineral association would indicate that themagmatic fluids interacted at a low fluid/rock ratio with thehost rock and/or mixed with reduced seawater-derived fluids.In both oxidized and reduced sequences the redox state of theinvolved fluids converges at the rock buffer line during coolingas a result of wall-rock reaction and fluid mixing, thus explain-ing why the later lower temperature pyrite-chalcopyrite suc-cession is shared by both mineral associations (Figs. 6,14, 15).

ConclusionsOre paragenesis and zoning, sulfur isotopes, thermody-

namic constraints, and alteration mineralogy and geochem-istry suggest that the Ral-Condestable deposit was formedby cooling and wall-rock buffering of metal-bearing magmaticfluids that, mainly in the upper parts of the system, mixedwith varying proportions of seawater-derived fluids whichcontributed a subordinate part of the sulfur.

The absence of rocks which could have oxidized hot (>300C)fluids to the field of hematite stability implies that the oxidizedmineral sequence (hematite-magnetite-pyrite-chalcopyrite)


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FeCl4--

hm

mtpy

po

0

-8

-7

-6

-5

-4

-3

-2

-1

R =

log

(fH/

fHO

)H

2(g

)2

(g)

0 50 100 150 200 250 300

T [C]

SO4

--

H S2(aq)

HSO4

-

H S2 (aq)

SO

4--

HS

O4-

pH =

5

pH =

4.5

high w

/r

low w/r

and/or mixing

rock buffer

pH = 4.5

pH

= 4

.5a S = 10

a Cl = 2

a Fe = 0.1

species-4

-

++

sulfur buffer

FIG. 15. RH-T diagram at a fixed fluid composition. Thermodynamic cal-culations done with the program Geochemists Workbench using the in-cluded thermo.com.v8.r6+.dat extended database. The diagram has beendrawn at a pH of 4.5 (solid line) and 5 (dotted line). Sulfur, Fe, and Cl activ-ities were adjusted so as to show the phase relationships that exist betweeniron oxides, sulfides, and Fe chlorides at temperatures below the upper300C limit admitted by the thermodynamic database. Although these activ-ities are not realistic at the higher temperatures considered for the main orestage (300400C, see text), extrapolation of the obtained pattern to highertemperatures agrees with the observed paragenetic relationships (Fig. 6).The triple point pyrite-pyrrhotite-magnetite moves to lower RH and highertemperature (see Fig. 14) at higher sulfur activity. In this diagram, the rockbuffer has been shifted from an RH of about 2.8 to 1.5. This value is notrealistic for a rock buffer but schematically shows how the high-temperaturereduced ore mineral sequence formed.

found in or near feeder structures precipited from magmaticbrines following the SO2-H2S gas buffer at high temperature(>350C) and fluid/rock ratio. This is supported by 34S ofsulfides from deep parts of feeder veins showing values rang-ing from 1.0 to 6.3 per mil (avg around 3.5), close to mag-matic. Furthermore, precipitation from externally derivedfluids could not explain the occurrence of the oxidized min-eral sequence, because these fluids would have been previ-ously reduced on their heating path by the basaltic to basalt-andesitic host sequence. Finally, mass balance based on sulfurisotope data suggests that at the deposit scale, the bulk of thesulfides is dominated by magmatic sulfur (probably around70%). The coeval tonalitic intrusion present in the core of thesystem is considered as the source of the mineralizing mag-matic fluids that also generated the poorly developed and bar-ren quartz stockwork and the associated potassic alteration.

High 34S values of sulfides ranging up to 26.3 per mil arefound in what is interpreted to have been a relatively shallowaquifer filled with reduced seawater. Heavy H2S was pro-duced through thermochemical reduction of Aptian seawatersulfate (34S = 14) in what corresponds to the hematite-chlorite (albite, epidote, calcite) alteration zone present at theupper flanks of the hydrothermal system, adjacent to thecausative intrusion. The seawater-derived fluids reachednear-chemical equilibrium with their host rock in the deeperactinolitized aquifer, at about 300 to 350C. Mixing of thesereduced fluids with magmatic brines already partially or to-tally reduced through reaction with wall rock at moderate tolow fluid/rock ratio can explain the large 34S scatter to heavyvalues observed in sulfides of the reduced mineral association(pyrrhotite-pyrite-chalcopyrite).

In terms of exploration criteria, and in areas devoid of oxi-dized rocks, the occurrence of early mushketovite (magnetitepseudomorphous after specular hematite) paragenetically fol-lowed by iron-bearing sulfides, as found in many IOCG de-posits, can be used as a direct field evidence for precipitationfrom oxidized magmatic brines. This observation is not onlyvalid for IOCG deposits, the same sequence being well docu-mented in skarns and porphyry deposits. The shallow subvol-canic position of the Ral-Condestable IOCG deposit and itsassociation with intermediate magmatism (Ral-Condestablesuperunit) are features commonly described for porphyrycopper and copper-skarn deposits. Sillitoe (2003) previouslypointed out that IOCG deposits can grade to Cu skarn min-eralization (e.g., parts of La Candelaria, Chile; Marschik andFontbot, 2001), and that several Cu-Au porphyry depositscontain abundant magnetite hematite and have alterationsimilar to IOCG deposits (abundant amphibole, clinopyrox-ene, sodic plagioclase, and magnetite). Although the presentstudy shows the importance of magmatic brines in the gene-sis of the Ral-Condestable deposit, the key parameters thatmake shallow (2- to 4-km depth) intermediate intrusion-cen-tered hydrothermal systems evolve to copper skarn, porphyrycopper, or IOCG deposits remain to be determined.

AcknowledgmentsThe present investigation was carried out with the support

of Ca. Minera Condestable S.A. and the Swiss National ScienceFoundation (grant 67836.02). We would like to acknowledgethe generous help provided by the Ca. Minera Condestable

staff, and in particular Juan Carlos Ortiz, Adalberto Riva -deneira, Patrick Dalla Valle, Carlos Rodriguez, Julio Ziga,Marco Carpio, Walter Cruz, and Emilio Cazimirio. Specialthanks are due to Yves Haeberlin and Jorge Spangenberg fortheir help in the sulfur isotope laboratory and to CatherineGinibre for the microprobe analyses at the University of Lau-sanne, Switzerland. We have also to acknowledge ChristophHeinrich, Fernando Corfu, Mario Sartori, Robert Moritz,Philip Schutte, Kalin Kouzmanov, and Dick Sillitoe for helpfulcomments and information, and Rossana Martini, Fabbio Cap-poni and Jean-Marie Boccard for technical help. The quality ofthis manuscript has been significantly improved thanks to thecareful work of the Economic Geology reviewers Pat Williams,Kurt Friehauf, and anonymous; their efforts are sincerely ac-knowledged. Special thanks are also due to Mark Hanningtonand Larry Meinert for their dedicated editorial work.February 15, 2007; February 24, 2009

REFERENCESAtkin, B.P., Injoque, J., and Harvey, P.K., 1985, Cu-Fe amphibole mineral-

ization in the Arequipa segment, in Pitcher, W.S., Atherton, M.P., Cobbing,E.J., and Beckinsale, R.D., eds., Magmatism at the plate edge: The Peru-vian Andes: Glasgow, UK, Blackie, p. 261270.

Barton, M.D., and Johnson, D.A., 1996, Evaporitic-source model for igneous-related Fe oxide-(REE-Cu-Au-U) mineralization: Geology, v. 24, p. 259262.

2000, Alternative brine sources for Fe-oxide (-Cu-Au) systems: Implica-tions for hydrothermal alteration and metals, in Porter, T.M., ed., Hy-drothermal iron oxide copper-gold and related deposits: A global perspec-tive: Adelaide, Australia, Australian Mineral Foundation, p. 4360.

Benavides, J., Kyser, T.K., Clark, A.H., Oates, C.J., Zamora, R., Tarnovschi,R., and Castillo, B., 2007, The Mantoverde iron oxide-copper-gold district,III Regin, Chile: The role of regionally derived, nonmagmatic fluids inchalcopyrite mineralization: ECONOMIC GEOLOGY, v. 102, p. 415440.

Boni, M., Rankin, A.H., and Salvadori, M., 1990, Fluid inclusion evidencesfor the development of Zn-Pb-Cu-F skarn mineralization in SW Sardinia,Italy: Mineralogical Magazine, v. 54, p. 279287.

Cardozo, M., 1983, Ral als Beispiel vulkanogener Kupferlagersttten imCopar-Metallotekt, Zentralperu: Unpublished Ph.D. thesis, Heidelberg,Ruprecht-Karls Universitt, 240 p.

Carten, R.B., 1986, Sodium-calcium metasomatism: Chemical, temporal,and spacial relationships at the Yerington, Nevada, porphyry copper de-posit: ECONOMIC GEOLOGY, v. 81, p. 14951519.

Chiaradia, M., Banks, D., Cliff, R., Marschik, R., and de Haller A., 2006, Ori-gin of fluids in iron oxide-copper-gold deposits: Constraints from 37Cl,87Sr/86Sri, and Cl/Br: Mineralium Deposita, v. 41, p. 565573.

Ciobanu, C.L., and Cook, N.J., 2004, Skarn textures and case study: the Ocnade Fier-Dognecea orefield, Banat, Romania: Ore Geology Reviews, v. 24,p. 315370.

Claypool, G.E., Holser, W.T., Kaplan, I.R., Sakai, H., and Zak, I., 1980, Theage curves of sulfur and oxygen isotopes in marine sulfate and their mutualinterpretation: Chemical Geology, v. 28, p. 199260.

de Haller, A., 2000, The Ral-Condestable iron oxides-Cu-Au deposit, Limadepartment, Peru: Preliminary results: Lule University of Technology Re-search Report 2000:6, p. 811.

2006, The Ral-Condestable iron oxide copper-gold deposit, centralcoast of Peru: Published Ph.D. thesis, Switzerland, University of Geneva,Terre et Environnement, v. 58, 123 p.

de Haller, A., Ziga, A. J., and Fontbot, L., 2001, Geology of the Ral-Con-destable iron oxide-copper-gold deposit, central coast of Peru [abs.]: Geo-logical Society of America Program with Abstracts, November 58, p. A129.

de Haller, A., Ziga, A. J., Corfu, F., and Fontbot, L., 2002, The iron oxide-Cu-Au deposit of Ral-Condestable, Mala, Lima, Peru [abs.]: CongresoGeolgico Peruano, 11th, Resmenes, p. 80.

de Haller, A., Corfu, F., Fontbot, L., Schaltegger, U., Barra, F., Chiaradia,M., Frank, M., and Ziga A., J., 2006, Geology, geochronology, and Hf andPb isotopic data of the Ral-Condestable iron oxide copper-gold deposit,central coast of Peru: ECONOMIC GEOLOGY, v. 101, p. 281310.

Dnkel, I., 2002, The genesis of East Elba iron ore deposits and their inter-relation with Messinian tectonics: Tbinger Geowissenschaftliche Arbeiten(TGA), A, Band 65, 143 p. + Appendix.


0361-0128/98/000/000-00 $6.00 383

Einaudi, M.T., Hedenquist, J.W., and Inan, E.E., 2003, Sulfidation state offluids in active and extinct hydrothermal systems: Transitions from por-phyry to epithermal environments: Society of Economic Geologists SpecialPublication 10, p.285313.

Fournier, R.O., 1999, Hydrothermal processes related to movement of fluidfrom plastic into brittle rock in the magmatic-epithermal environment:ECONOMIC GEOLOGY, v. 94, p. 11931211.

Frikken, P.H., Cooke, D.R., Walshe, J.L., Archibald, D., Skarmeta, J., Ser-rano, L., and Vargas, R., 2005, Mineralogical and isotopic zonation in theSur-Sur tourmaline breccia, Ro Blanco-Los Bronces Cu-Mo deposit, Chile:Implications for ore genesis: ECONOMIC GEOLOGY, v. 100, p. 935961.

Giesemann, A., Jager, H.J., Norman, A.L., Krouse, H.P., and Brand, W.A.,1994, Online sulfur-isotope determination using an elemental analyzer cou-pled to a mass-spectrometer: Analytical Chemistry, v. 66, p. 28162819.

Giggenbach, W.F., 1984, Mass transfer in hydrothermal alteration systemsaconceptual approach: Geochimica et Cosmochimica Acta, v. 48, p. 26932711.

1987, Redox processes governing the chemistry of fumarolic gas dischargesfrom White Island, New Zealand: Applied Geochemistry, v. 2, p. 143161.

1992, SEG Distinguished lecture: Magma degassing and mineral depo-sition in hydrothermal systems along convergent plate boundaries: ECO-NOMIC GEOLOGY, v. 87, p. 19271944.

1997, The origin and evolution of fluids in magmatic-hydrothermal sys-tems, in Barnes, H.L., ed., Geochemistry of hydrothermal ore deposits, 3rded.: New York, NY, J. Wiley and Sons, p. 737796.

Haynes, D.W., 2000, Iron oxide copper (-gold) deposits: Their position in theore deposit spectrum and modes of origin, in Porter, T.M., ed., Hydrother-mal iron oxide copper-gold and related deposits: A global perspective: Ade-laide, Australia, Australian Mineral Foundation, p. 4360.

Helgeson, H.C., Delany, J.M., Nesbitt, H.W., and Bird, D.K., 1978, Sum-mary and critique of the thermodynamic properties of rock-forming min-erals: American Journal of Science, v. 278a, 229 p.

Henley, R.W., and Ellis, A.J., 1983, Geothermal systems ancient and modern:A geochemical review: Earth-Science Reviews, v. 19, p. 150.

Hitzman, M.W., 2000, Iron oxide-Cu-Au deposits: What, where, when, andwhy, in Porter, T.M., ed., Hydrothermal iron oxide copper-gold and relateddeposits: A global perspective: Adelaide, Australia, Australian MineralFoundation, p. 4360.

Huston, D.L., 1999, Stable isotopes and their significance for understandingthe genesis of volcanic-hosted massive sulfide deposits: a review: Reviewsin Economic Geology, v. 8, p. 157179.

Injoque, J., 1985, Geochemistry of the Cu-Fe-amphibole skarn deposits ofthe Peruvian central coast: Unpublished Ph.D. thesis, UK, University ofNottingham, 359 p.

2002, Fe oxide-Cu-Au deposits in Peru: An integrated view, in Porter, T.M., ed., Hydrothermal iron oxide copper-gold and related deposits: Aglobal perspective: Adelaide, Australia, Porter GeoConsultancy Publishing,v. 2, p. 97113.

Leake, B.E., Wooley, A.R., Arps, C.E.S., Birch, W.D., Gilbert, M.C., Grice,J.D., Hawthorne, F.C., Kato, A., Kisch, H.J., Krivovichev, V., Linthout, K.,Laird, J., Mandarino, J.A., Maresch, W.V., Nickel, E.H., Rock, N.M.S.,Schumacher, J.C., Smith, D.C., Stephenson, N.C.N., Ungaretti, L., Whit-taker, E.J.W., and Youzhi, G., 1997, Nomenclature of amphiboles: Reportof the subcommittee on amphiboles of the International Mineralogical As-sociation, commission on new minerals and mineral names: Canadian Min-eralogist, v. 35, p. 219246.

Marschik R., and Fontbot L., 2001, The Candelaria-Punta del Cobre iron oxideCu-Au (-Zn-Ag) deposits, Chile: ECONOMIC GEOLOGY, v. 96, p. 17991826.

Marschik, R., and Sllner, F., 2006, Early Cretaceous U-Pb zircon ages for theCopiapo plutonic complex and implications for the IOCG mineralization atCandelaria, Atacama region, Chile: Mineralium Deposita, v. 41, p. 785801.

Mcke, A., and Cabral, A.R., 2005, Redox and nonredox reactions of mag-netite and hematite in rocks: Chemie der Erde, v. 65, p. 271278.

Myers, J., and Eugster, H.P., 1983, The system Fe-Si-O: Oxygen buffer calibra-tions to 1,500K: Contributions to Mineralogy and Petrology, v. 82, p. 7590.

Ohmoto, H., 1972, Systematics of sulfur and carbon isotopes in hydrothermalore deposits: ECONOMIC GEOLOGY, v. 67, p. 551578.

1986, Stable isotope geochemistry of ore deposits: Reviews in Mineral-ogy, v. 16, p. 491559.

1996, Formation of volcanogenic massive sulfide deposits: the Kurokoperspective: Ore Geology Reviews, v. 10, p.135177.

2003, Nonredox transformation of magnetite-hematite in hydrothermalsystems: ECONOMIC GEOLOGY, v. 98, p. 157161.

Ohmoto, H., and Goldhaber, M.B., 1997, Sulfur and carbon isotopes, inBarnes, H.L., ed., Geochemistry of hydrothermal ore d 3rd ed.: New York,J. Wiley and Sons, p. 517611.

Ohmoto, H., and Rye, R.O., 1979, Isotope of sulfur and carbon, in Barnes,H.L., ed., Geochemistry of hydrothermal ore deposits, 2nd ed.: New York,J. Wiley and Sons, p. 509567.

Otake, T., Wesolowski, D.J., Anowitz, L.M., Allard, L.F., and Ohmoto, H.,2007, Experimental evidence for non-redox transformations between mag-netite and hematite under H2-rich hydrothermal conditions: Earth andPlanetary Science Letters, v. 257, p. 6070.

Palacios, O., Caldas, J., and Vela, C., 1992, Geologa de los cuadrngulos deLima, Lurn, Chancay y Chosica: Lima, Peru, INGEMMET Boletn 43,Serie A, Carta Geolgica Nacional, 163 p.

Pollard, P.J., 2000, Evidence of a magmatic fluid and metal source for Fe-oxide Cu-Au mineralization, in Porter, T.M., ed., Hydrothermal iron oxidecopper-gold and related deposits: A global perspective: Adelaide, Australia,Australian Mineral Foundation, p. 2741.

2001, Sodic(-calcic) alteration in Fe-oxide-Cu-Au districts: An origin viaunmixing of magmatic H2O-CO2-NaCl CaCl2-KCl fluids: MineraliumDeposita, v. 36, p. 93100.

Prendergast, K., Clarke, G.W., Pearson, N.J., and Harris, K., 2005, Genesisof pyrite-Au-As-Zn-Bi-Te zones associated with Cu-Au skarns: Evidencefrom the Big Gossan and Wanagon gold deposits, Ertsberg district, Papua,Indonesia: ECONOMIC GEOLOGY, v. 100, p. 10211050.

Ramdohr, P., 1980, The ore minerals and their intergrowths: Oxford, UK,Pergamon Press, 1,205 p.

Reed, M.H., 1997, Hydrothermal alteration and its relationship to ore fluidcomposition, in Barnes, H.L., ed., Geochemistry of hydrothermal ore de-posits, 3rd ed.: New York, J. Wiley and Sons, p. 303365.

Ripley, E.M., and Ohmoto, H., 1977, Mineralogic, sulfur isotope and fluid in-clusion studies of the stratabound copper deposits at Raul mine, Peru:ECONOMIC GEOLOGY, v. 72, p. 10171041.

Ripley, E.M., and Ohmoto, H., 1979, Oxygen and hydrogen isotopic studiesof ore deposition and metamorphism at the Raul mine, Peru: Geochimicaet Cosmochimica Acta, v. 43, p. 16331643.

Rivera, R., Petersen, G., and Rivera, M., 1975, Estratigrafa de la costa deLima: Boletn de la Sociedad Geolgica del Per, v. 45, p. 159186.

Salazar, H., and Landa, C., 1993, Geologa de los cuadrngulos de Mala, Lu-nahuan, Tupe, Conayca, Chincha, Tantar y Castrovirreyna: Lima, Peru,INGEMMET Boletn 44, Serie A: Carta Geolgica Nacional, 91 p.

Shanks, W.C., III, 2001, Stable isotopes in seafloor hydrothermal systems:Vent fluids, hydrothermal deposits, hydrothermal alteration, and microbialprocesses: Reviews in Mineralogy and Geochemistry, v. 43, p. 469525.

Shanks, W.C., III, and Seyfried, W.E., Jr., 1987, Stable isotope studies of ventfluids and chimney minerals, southern Juan de Fuca Ridge: Sodium meta-somatism and seawater sulfate reduction: Journal of Geophysical Research,v. 2, p. 11,38711,399.

Shanks, W.C., III, Bischoff, J.L., and Rosenbauer, R.J., 1981, Seawater sul-fate reduction and sulfur isotope fractionation in basaltic systems: Interac-tion of seawater with fayalite and magnetite at 200350C: Geochimica etCosmochimica Acta, v. 45, p. 19771995.

Sillitoe, R.H., 2003, Iron oxide-copper-gold deposits: An Andean view: Min-eralium Deposita, v. 38, p. 787812.

Vaughan, D.J., and Craig, J.R., 1997, Sulfide ore mineral stabilities, mor-phologies, and intergrowth textures, in Barnes, H.L., ed., Geochemistry ofhydrothermal ore deposits, 3rd ed.: New York, J. Wiley and Sons, p. 367434.

Vidal, C., Injoque, J., Sidder, G.B., and Mukasa, S.B., 1990, Amphibolitic Cu-Fe skarn deposits in the Central Coast of Peru: ECONOMIC GEOLOGY, v. 85,p. 14471461.

Wauschkuhn, A., 1979, Geology, mineralogy and geochemistry of thestratabound Cu deposits in the Mesozoic Coastal belt of Peru: Boletn de laSociedad Geolgica del Per, v. 62, p. 161192.

Whitney, J.A., Hemley, J.J., and Simon, F.O., 1985, The concentration of ironin chloride solutions equilibrated with synthetic granitic compositions: thesulfur-free system: ECONOMIC GEOLOGY, v. 80, p. 444460.

Williams, P.J., Barton, M.D., Johnson, D.A., Fontbot, L., de Haller, A.,Mark, G., Oliver, N.H.S., and Marschik, R., 2005, Iron oxide copper-golddeposits: Geology, space-time distribution, and possible modes of origin:ECONOMIC GEOLOGY 100TH ANNIVERSARY VOLUME, p. 371405.

Zheng, Y.-F., and Hoefs, J., 1993, Effects of mineral precipitation on the sul-fur isotope composition of hydrothermal solutions: Chemical Geology, v.105, p. 259269.


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