Stable isotopic constraints on Kuroko-type paleohydrothermal systems

in the Mesoproterozoic Serra do Itaberaba group, Sao Paulo State, Brazil

Annabel Perez-Aguilara,*, Caetano Juliania, Lena V.S. Monteirob,Anthony E. Fallickc, Jorge S. Bettencourta

aInstituto de Geociencias, Universidade de Sao Paulo, Rua do Lago, 562, CEP 05508-080, Sao Paulo, SP, BrazilbInstituto de Geociencias, Universidade Estadual de Campinas, Rua Joao Pandia Calogeras, 51, CEP 13083-970, Campinas, SP, Brazil

cScottish Universities Environmental Research Centre, Rankine Avenue, East Kilbride, Glasgow G75 OQF, Scotland, UK

Accepted 1 November 2004

Abstract

Mesoproterozoic oceanic paleohydrothermal systems developed in the volcanosedimentary Serra do Itaberaba Group, which comprises

part of the Ribeira fold belt. Hydrothermal alteration associated with these systems was responsible for large premetamorphic chloritic

alteration halos (CZ1 rocks), overprinted by restricted premetamorphic chloritic (CZ2 rocks), argillic, and advanced argillic alterations that

correspond to intensely leached rocks within feeder zones. Well-defined trends of increasing d18O values with the progressive intensity of the

alteration process are observed for igneous metabasites, metabasic hydroclastic rocks, and intermediate metamorphosed igneous and

volcaniclastic rocks from CZ1. Systematic stable isotope variations evince that, in the Serra do Itaberaba metamorphosed hydrothermalized

rocks, the preexisting isotope signatures of the hydrothermal systems were at least partially preserved. Highly evolved hot seawater is

suggested for the genesis of the CZ1 rocks, whereas for the CZ2 rocks and marundites, the 18O fluid enrichments are interpreted as due to the

major contribution of evolved seawater-derived fluids with a subordinate magmatic water component. An early near-seafloor, low-

temperature alteration in a mid-ocean ridge environment was responsible for heterogeneous 18O whole-rock enrichments and followed by

steady hydrothermal circulation with discharge of hot fluids, which previously underwent isotopic exchange with the 18O enriched volcanic

rocks in the deeper part of the system with high temperatures and low water: rock ratios in a backarc environment. The subordinate magmatic

water component derived from andesitic and rhyodacitic intrusions. The extremely high d18O anomalies from the CZ1 rocks suggest an

associated base metal massive sulfide ore body. The lower d18O values related to the CZ2 rocks represent alteration by a higher temperature

fluid, which might indicate the proximity of possible ore zones. The identification of several premetamorphic hydrothermally altered zones,

similar to those of Kuroko-type base metal mineralizations, expands the mineral potential of base metal deposits in the Serra do Itaberaba

Group and the volcanosedimentary sequences from the Ribeira fold belt.

q 2005 Elsevier Ltd. All rights reserved.

Keywords: Kuroko-type deposits; Mesoproterozoic; Paleohydrothermal system; Serra do Itaberaba group; Stable isotopes; VMSD

1. Introduction

Cummingtonite-anthophyllite-cordierite-garnet rocks

from metamorphosed volcanosedimentary sequences

usually are interpreted as metamorphic products of hydro-

thermalized basic to acid volcanic rocks (James et al., 1978;

Riverin and Hodgson, 1980; Spear, 1982; Elliott-Meadows

and Appleyard, 1991; Roberts et al., 2003). These rocks,

0895-9811/$ - see front matter q 2005 Elsevier Ltd. All rights reserved.

doi:10.1016/j.jsames.2004.11.012

* Corresponding author. Tel.: C55 11 50772160; fax: C55 11 50772219.

E-mail address: anaperez99@hotmail.com (A. Perez-Aguilar).

formed under a medium to high metamorphic grade, often

are affected by intense tectonic transposition, which makes

the reconstitution of the geometry of the hydrothermal

alteration zones difficult.

Rocks formed in paleohydrothermal systems have been

identified in the Mesoproterozoic, medium-grade metamor-

phosed Serra do Itaberaba Group (Juliani et al., 1986, 2000a,

2000b; Perez-Aguilar et al., 2000, 2002a, 2002b). Despite

metamorphism and deformation, well-defined zones gener-

ated by different types and intensities of hydrothermal

alteration remain recognizable in these rocks. The alteration

zones encompass metamorphic rocks composed of (1)

cummingtonite, anthophyllite, gedrite, cordierite, garnet,

Journal of South American Earth Sciences 18 (2005) 305–321

www.elsevier.com/locate/jsames

A. Perez-Aguilar et al. / Journal of South American Earth Sciences 18 (2005) 305–321306

Mg-chlorite, staurolite, ilmenite, rutile, Ca-rich plagioclase,

and quartz; (2) Mg-chlorite, cummingtonite, garnet,

magnesiohornblende, and tschermakite; (3) corundum,

margarite, and rutile; (4) diopside, actinolite, epidote,

carbonate, plagioclase, and quartz; (5) biotite in biotite-

rich metabasites and metamorphosed intermediate rocks;

and (6) hydrothermal quartz in quartz-rich metabasites and

metamorphosed intermediate rocks (Juliani et al., 1994;

Perez-Aguilar, 1996, 2001; Perez-Aguilar et al., 2000,

2002b). These metamorphic rocks show exotic bulk

chemical compositions and mineralogical association that

are attributable to the intense premetamorphic leaching and

metasomatism of basic to acid igneous and volcaniclastic

protoliths affected by chloritic, argillic, and advanced

argillic alterations, carbonatization, potassification, and

silicification. These alteration processes are similar to

those present in Kuroko-type base metal deposits (Sangster,

1972; Ishihara et al., 1974; Franklin, 1993; Ohmoto, 1996;

Shikazono, 2003).

Stable isotope data can be fundamental for the determi-

nation of the fluid sources, temperature, and evolution of

hydrothermal systems and thereby constrain genetic models.

In addition, the application of stable isotopes in mineral

exploration has been emphasized in the past two decades,

mainly because of the recognition of isotope halos as ore

guides (Beaty and Taylor, 1982; Criss and Taylor, 1983;

Beaty et al., 1988; Cathles, 1993; Waring et al., 1998).

However, regional metamorphic events and later hydrother-

mal overprints may obliterate original isotope compositions

and cause either isotopic homogenization or shifts.

Fig. 1. Regional geological map with the locatio

The oxygen and hydrogen isotopic variations observed in

the hydrothermalized and metamorphosed rocks of the Serra

do Itaberaba Group indicate that preexisting isotope

gradients were preserved, at least partially, in a way similar

to well-documented examples of metamorphosed Precam-

brian massive base metal deposits (Beaty and Taylor, 1982;

Beaty et al., 1988; Araujo et al., 1996). Thus, a study based

on oxygen stable isotope halos could identify exploration

targets and explicate the metallogenetic potential of the

Mesoproterozoic volcanosedimentary sequences in south-

east Brazil.

2. Geological setting

Paleohydrothermal systems are located northeast of Sao

Paulo (Fig. 1) and hosted by the Mesoproterozoic Serra do

Itaberaba Group (Fig. 2), which is partially covered by the

Neoproterozoic Sao Roque Group (Juliani et al., 1986,

2000a, 2000b; Hackspacher et al., 1999). Both are part of

the Ribeira fold belt (Almeida et al., 1973). The whole

sequence was intruded by several Neoproterozoic-

Phanerozoic granitic plutons and affected by several NE-

SW–trending shear zones (Almeida et al., 1981).

The Serra do Itaberaba Group (Fig. 2), which is comprised

of the Morro da Pedra Preta, Nhangucu, and Pirucaia

Formations (Juliani, 1993; Juliani and Beljavskis, 1995),

was affected by two progressive regional metamorphic

events that record clockwise P-T-t paths (Juliani et al.,

1997, 2000a). The first, Mesoproterozoic metamorphic event

n of the study area (Juliani et al., 2000a).

Fig. 2. Geological map of Itaberaba and Pedra Branca hills showing the location of the paleohydrothermal systems (Juliani, 1993; Juliani et al., 2000a).

A.

Perez-A

gu

ilar

eta

l./

Jou

rna

lo

fS

ou

thA

merica

nE

arth

Scien

ces1

8(2

00

5)

30

5–

32

13

07

Fig. 3. Schematic reconstruction of the hydrothermal system.

A. Perez-Aguilar et al. / Journal of South American Earth Sciences 18 (2005) 305–321308

ranges from the Barrovian upper greenschist to the upper

amphibolite facies (490–650 8C, 4–7 kbar). The second,

Neoproterozoic metamorphic event affected the Serra do

Itaberaba Group in the upper greenshist to the amphibolite

facies (500–580 8C) in a low-pressure regime (4–4.7 kbar).

The final metamorphic evolution is given by retrometa-

morphism in the greenschist facies (Juliani et al., 1997,

2000a).

The basal Morro da Pedra Preta Formation is composed

of a metamorphosed normal mid-ocean ridge basalt

(N-MORB) with pillow lavas, basic volcanic agglomerates,

breccias, lapillistones, lapilli tuffs, and tuffs. It is covered by

metapelites, graphite schists, sulfide-rich schists, and

manganiferous schists, with subordinate metabasalts, basic

to acid metatuffs, Algoma-type banded-iron formations

(BIFs), tourmalinites, and calc-silicate rocks. In the upper

metabasalts of the Morro da Pedra Preta Formation, small

dome-like brecciated intrusions of andesite, dacite, and

rhyodacite occur, surrounded by volcanic breccias and tuffs

(Juliani, 1993). The paleohydrothermal systems are

spatially related to these intrusions. The overlying Nhan-

gucu Formation comprises iron-manganiferous schists,

calc-silicate schists, and small lenses of metabasalts,

metatuffs, and marbles covered by andalusite-chlorite

schists. The Nhangucu Formation was generated in a

backarc basin produced by a westward ensimatic subduc-

tion. The andesitic, dacitic, and rhyodacitic intrusions are

mainly related to this subduction event. The Pirucaia

Formation comprises quartzites and quartz-rich schists,

which represent a shoreline sedimentary facies of the

Nhangucu Formation.

3. Serra do Itaberaba paleohydrothermal systems

The Serra do Itaberaba paleohydrothermal systems are

spatially and genetically linked to the andesitic-rhyodacitic

shallow intrusions (Figs. 3 and 4) emplaced during backarc

basin evolution (Juliani et al., 1992; Perez-Aguilar, 1996,

2001). Large metamorphosed chloritic alteration zones

(CZ1), similar to those described in metamorphosed

volcanogenic massive sulfide deposits (VMSDs) (Riverin

and Hodgson, 1980; Elliott-Meadows and Appleyard, 1991;

Roberts et al., 2003), surround hydrothermal feeder zones.

Within hydrothermal feeder zones, premetamorphic chlori-

tic (CZ2), argillic, and advanced argillic alteration occur

(Fig. 3). A diffuse zone of K-enrichment, marked by biotite-

bearing basic and intermediate rocks, envelops the CZ1,

which defines a lower temperature potassic alteration.

Premetamorphic carbonatization zones typically occur in

the interface of basaltic flows in deeper parts of the system

but also along fracture-controlled hydrothermal channel-

ways. Potassic alteration and silicification overprint the

early hydrothermalized rocks. Algoma-type BIFs, sulfide-

rich metapelites, and gold mineralizations are also geneti-

cally and spatially related to the hydrothermal center

(Juliani, 1993; Perez-Aguilar, 1996). The schematic recon-

struction of the hydrothermal zones appears in Fig. 3. The

geometry of the hydrothermally altered rocks broadly

includes inverted cone shapes that flare upward (Fig. 4),

despite the overprinting of intense deformation processes.

The premetamorphic CZ1 event affected basic, inter-

mediate, and acid igneous and volcaniclastic rocks with

variable intensity. Complete gradation in the metamorphic

products includes weakly, transitional, moderately, and

strongly altered rocks from the outermost zone to the inner

part of the alteration zone (Fig. 5; Perez-Aguilar, 1996,

2001). The metamorphic products of chloritized rocks from

the CZ1 are recognized by the presence of variable amounts

of anthophyllite, gedrite, and/or cummingtonite (Fig. 6).

Despite the different original compositions of altered rocks

from the CZ1, the weakly altered rocks can be identified by

the presence of small amounts of cordierite and/or

cummingtonite. Transitional rocks typically have two or

more coexistent amphiboles (magnesiohornblende, tscher-

makite, anthophyllite, cummingtonite, or gedrite). Moder-

ately altered rocks show total replacement of hornblende by

cummingtonite and small amounts of cordierite. In addition,

two alteration zones can be distinguished: garnet-free (PZ1)

and garnet-bearing (PZ2). Strongly altered rocks derived

from felsic and mafic protoliths have a similar metamorphic

mineralogical composition that matches the alteration

patterns related to cordierite-anthophyllite rocks from

Tunaberg, Sweden (Dobbe, 1994); Manitouwadge, Canada

(Pan and Fleet, 1995); and Ruostesuo, central Finland

(Roberts et al., 2003). These strongly altered rocks

typically have radiate clusters of coarse-grained cumming-

tonite, gedrite, and/or anthophyllite and Mg-cordierite

poikiloblasts (Fig. 7a), as well as variable amounts of

almandine poikiloblasts, quartz, magnetite, ilmenite, rutile,

staurolite, biotite, Mg-chlorite, phlogopite, and gedrite

(Perez-Aguilar, 1996, 2001; Perez-Aguilar et al., 2000).

The CZ2 metamorphic products are represented by

rocks composed of magnesiohornblendeCtschermakiteGMg-chlorite or Mg-chloriteGcummingtoniteGgarnetG

Fig. 4. Geological map of the paleohydrothermal system area (modified from Perez-Aguilar, 1996).

A.

Perez-A

gu

ilar

eta

l./

Jou

rna

lo

fS

ou

thA

merica

nE

arth

Scien

ces1

8(2

00

5)

30

5–

32

13

09

Fig. 5. Schematic representation of hydrothermalized lithotypes, showing

gradation between unaltered to strongly hydrothermally altered metabasic

and intermediated metamorphosed rocks. (1) Unaltered metabasic rocks,

(2) weakly altered metabasic rocks, (3) metabasic rocks from the

transitional zone, (4) moderately altered metabasic rocks, (5) strongly

altered metabasic rocks, (6) strongly altered intermediated metamorphosed

rocks, (7) moderately altered meta-intermediate rocks, (8) intermediated

metamorphosed rocks from the transitional zone, and (9) weakly altered

and unaltered intermediated metamorphosed rocks. For scale, see Fig. 3.

A. Perez-Aguilar et al. / Journal of South American Earth Sciences 18 (2005) 305–321310

plagioclase; we refer to them as metachloritites. The rocks

display higher Mg enrichments and Si depletions in

bulk chemical composition compared with rocks

from the CZ1. Field relationships show that the CZ1 is

crosscut by the CZ2, and veins composed essentially of

magnesiohornblendeGquartzGgarnet, which cut strongly

altered CZ1 rocks, could be related to the CZ2 event.

The original geometries of CZ1 and CZ2 (Fig. 4) are

similar to those typically present beneath massive sulfide

mineralizations, which have been interpreted as chloritie

hydrothermal pipes (Sangster, 1972; Spence and de Rosen-

Spence, 1975; Schermerhorn, 1978; MacGeehan and

MacLean, 1980; Beaty et al., 1982).

Metatuffs from the CZ1 grade continuously to quartz-free

sericite and/or muscovite schists, then to margarite and/or

muscoviteCcorundumGrutileGCa-plagioclaseGtourma-

line rocks, and finally to muscoviteCcorundumGrutile rocks

Fig. 6. Outcrop of intermediate-composition metavolcaniclastic hydro-

thermalized rocks, showing gradation between weakly altered to strongly

altered rocks. (1) Weakly altered rocks, (2) rocks from the transitional zone

with hornblende predominating over cummingtonite, (3) rocks from the

transitional zone with cummingtonite predominating over hornblende,

(4 PZ1) moderately altered garnet-free rocks, (4 PZ2) moderately altered

garnet-bearing rocks, and (5) strongly altered rocks. Dark lines represent

boundaries between different altered lithotypes.

Fig. 7. (A) Photomicrograph of a strongly altered rock of original

intermediate composition: (1) garnet poikiloblast, (2) cummingtonite, and

(3) cordierite poikiloblast; trasmitted light, crossed polars, wide side of

photo, 5.5 mm. (B) Meta volcaniclastic (1) weakly altered rock, (2) rock

from the transitional zone, and (3) rock from the transitional zone

overprinted by potassic alteration. (C) Photomicrograph of a strongly

altered rock of original intermediate composition: (1) garnet-quartz

intergrowths in garnet poikiloblast, (2) Mg-amphibole, (3) cordierite

pokiloblast; wide side of photo, 5.5 mm; transmitted light.

or margariteCcorundumGrutile rocks. We refer to these

corundum-bearing rocks as marundites, the name used by Hall

(1920) for rocks formed of margariteCcorundumCrutile in

the Barberton greenstone belt. Chemical compositions of the

marundites from the Serra do Itaberaba Group are similar to

those of Al-rich clay rocks found in association with Kuroko-

type deposits (Schmidt, 1985; Shikazono, 2003). Thus, these

rocks have been interpreted as the metamorphic products

of argillic and advanced argillic alterations (Juliani, 1993;

A. Perez-Aguilar et al. / Journal of South American Earth Sciences 18 (2005) 305–321 311

Juliani et al., 1994). The metamorphic product of the

carbonatization, which occurs in the deeper part of the system,

includes diopsideCactinoliteCcarbonateCepidoteCplagio-

claseCquartz rocks.

4. Sample preparation and analytical methods

4.1. Sample preparation

Whole-rock and mineral (quartz, garnet, muscovite, and

margarite) separates for stable isotope analysis were

prepared from select samples after detailed petrographic,

paragenetic, and microstructural studies. The garnet poiki-

loblast (Fig. 7c) extraction from the rocks was first

performed using a circular silica carbide dental drill,

followed by fine crushing in an agate mortar to less than

200 mesh to eliminate fine garnet and quartz intergrowths.

Subsequently, these minerals were separated electromagne-

tically and by density using bromoform. Finally, the clean

grains were hand picked using a stereomicroscope and an

adapted needle. Despite this procedure, w1% of very fine-

grained opaque mineral inclusions, probably ilmenite,

remain in quartz and garnet. The garnet also shows small

amounts of very fine-grained quartz. Margarite and

muscovite were separated from monomineralic margarite

and muscovite schists with a metallic stick.

Whole-rock samples were crushed in a hydraulic press

using two tungsten carbide plates until the rock fragments

were !5 mm. The samples were pulverized until O97%

went through a 200 mesh, then homogenized and split. For

the final sample, in excess of 50 g was obtained for whole-

rock oxygen and hydrogen analyses. Whole-rock 18O/16O

ratios were measured with 37 samples, and whole-rock dD

was measured with 2 samples.

4.2. Analytical techniques

Isotopic analyses were conducted at Scottish Universities

Environmental Research Centre (SUERC). Whole-rock and

silicate 18O/16O analyses were made from w1 mg rock

powders or separated minerals using a laser fluorination

system, based on the method described by Sharp (1990).

Oxygen was released from the samples by heating them

with a CO2 laser inside a ClF3-charged chamber. Sub-

sequently, the released oxygen was converted to CO2 and

analyzed on a VG PRISM III mass spectrometer. In this

technique, no laser-correcting factor is required because

each sample is reacted to completion, so all oxygen is

collected. The analytical precision of laser fluorination is

0.2‰ at one sigma. The NSB 30 biotite gives 5.4‰.

The hydrogen isotope analyses were performed on

w40 mg whole-rock and mineral phases previously heated

to 120 8C overnight to remove absorbed volatiles (Fallick

et al., 1993). The samples then were dehydroxylated by

heating to 1400 8C. Water vapor and CO2 were collected

and cryogenetically separated, then the water vapor was

reduced to H2 in a chromium furnace at 830 8C (Donnelly

et al., 2001). The collected H2 was transferred into a

Micromass Optima mass spectrometer. The analytical

precision of this technique is approximately 0.2‰ at one

sigma, and the NSB 30 biotite gives K65‰. The oxygen

and hydrogen isotope results are expressed in conventional

delta (d) notation, per mil (‰), and relative to Vienna

standard mean ocean water values (V-SMOW).

5. Results

5.1. Whole-rock stable isotope compositions

Metabasites derived from igneous protoliths show d18O

values from C5.9 to C16.9‰ (Table 1, Fig. 8). The

unaltered metabasites, without typical hydrothermal-related

minerals, display d18O values from C5.9 to C9.0‰. The

d18O values in samples from the transitional zone are C8.6

to C10.8‰. In the strongly hydrothermally altered samples,

the d18O values range from C11.8 to C16.9‰ (Table 1,

Fig. 8). The oxygen isotope value of a metamorphosed,

unaltered, basic metalapillistone is C8.3‰, and that of the

metamorphic, hydrothermally altered lapilli-tuff from the

transitional zone is C10.1‰ (Table 1, Fig. 9).

Hydrothermally altered intermediate meta-igneous

rocks (Table 1) yield d18O values ranging from C14.1 to

C17.6‰. The lower d18O values (C14.1 to C15.4‰)

correspond to samples from the transitional zone, and the

higher d18O values (C16.8 to C17.6‰) correspond to

moderately and strongly hydrothermally altered rocks

(Table 1, Fig. 8).

The metavolcaniclastic rocks of basaltic andesitic and

andesitic composition show d18O values of C15.4‰ for the

unaltered protolith, C16.9‰ for weakly hydrothermally

altered rock, C15.3 to C17.4‰ for rocks from the

transitional zone, C16.4 to C17.0‰ for moderately

hydrothermally altered rocks, and C16.5 to C17.8‰ for

most strongly hydrothermally altered rocks. However, one

of the strongly altered samples (301a) has d18OZC11.6‰.

In this group of rocks, the following samples belong to a

continuous outcrop: 635b8 (weakly altered rock), 635c4

(transitional zone rock), 635c1 (moderately altered rock,

PZ2), and 635 h (strongly altered rock).

The intermediate meta-igneous rocks from the moder-

ately altered and transitional zone overprinted by a potassic

alteration (samples 440b and 417a) show values of d18O of

C15.9 to C16.2‰ (Table 1, Fig. 8). Two rock samples

from the CZ2 yield d18O values of C9.0 and C10.6‰, and

one displays whole-rock dDZK88‰. The plagioclase

marundite shows an d18O value of C9.7‰ and dD of

K55‰ (Table 1, Fig. 8).

Table 1

Mineralogical and stable isotope compositions of nonaltered and altered metabasites and intermediate meta-igneous and metavolcaniclastic rocks

Sample Type of analysis Lithology Mineralogical composition (% volume) d18O (‰) dD (‰)

Rocks from chloritic alteration zone 1 (CZ1)

194c WR MBI-NA Hbl(60) Pl(37) Op(3) 5.9 –

484a WR MBI-NA Hbl(60) Pl(35) Op(5) 7.7 –

260a WR MBI-NA Hbl(50) Pl(40) Qtz(5) Op(5) 7.8 –

123a WR MBI-NA Hbl(70) Pl(25) Op(3) Ttn(2) 9.0 –

2783b WR MBI-TZ Hbl(45) Cum(15) Pl(30) Qtz(5) Op(5) 8.6 –

352aa WR MBI-TZ MgHblCTs(23) CumCGed(50) Pl(20) Op(7) 10.8 –

211a WR MBI-SA Cum(50) Pl(35) Qtz(10) Op(5) 16.9 –

127bb,a WR MBI-SA Cum(60) AthCGed(5) Pl(28) Op(7) 11.8 –

V582d WR MBL-NA Hbl(80) Pl(13) Op(7) 8.3 –

580b WR MBT-TZ Hbl(45) CumCGed(25) Pl(23) Op(7) 10.1 –

507d1 WR MII-TZ Hbl(25) Cum(30) Qtz(28) Pl(15) Op(2) 14.6 –

507d2 WR MII-TZ Hbl(20) Cum(35) Qtz(30) Pl(13) Op(2) 14.1 –

177 WR MII-TZ Hbl(35) Cum(15) Qtz(20) Pl(10) Crd(15) Op(5) 15.4 –

177c 15.5 –

2756a WR MII-MA Cum(40) Grt(15) Qtz(20) Pl(12) Crd(5) Chl(5) St(3) 16.8 –

2756ac 16.5 –

110f WR MII-SA Cum(20) Ath(10) Crd(28) Qtz(25) Grt(10)

Op(5) Chl(2)

17.2 –

183 WR MII-SA Ath(15) Cum(8) Crd(46) Qtz(27) Op(4) 17.6 –

291c WR MIV-NA Hbl(67) Grt(10) Qtz(15) Pl(4) Op(2) 15.5 –

194b2 WR MIV-TZ Hbl(15) Cum(35) Qtz(23) Pl(17) Grt(5) Op(5) 16.2 –

296b WR MIV-TZ Hbl(38) Cum(27) Pl(15) Grt(7) Qtz(10) Op(3) 15.3 –

176e WR MIV-SA Cum(38) Crd(35) Qtz(10) Pl(10) Op(7) 16.5 –

121a WR MIV-TZ Hbl(35) Cum(20) Qtz(28) Pl(10) Grt(5) Op(2) 16.8 –

300b WR MIV-TZ Hbl(20) Cum(35) Qtz(33) Pl(12) Op(5) 17.2 -

1100 WR MIV-MA/PZ1 Cum(40) Crd(23) Qtz(21) Grt(9) Pl(5) St(2) 16.9 –

311a WR MIV-MA/PZ2 Cum(23) Ged(5) Qtz(30) Pin(15) Grt(7) Op(7) Pl(3)

Chl(5) Crd(5)

16.4 –

100 WR MIV-SA Ath(32) Cum(6) Qtz(30) Grt(16) Crd(11)

Op(3) Zo(2)

16.9 –

101a WR MIV-SA Ged(32) Qtz(27) Grt(19) Crd(17) Op(3) Ep(2) 17.2 –

301a WR MIV-SA Ath(25) Cum(10) Qtz(30) Crd(20) Pl(10) Op(5) 11.6 –

301ac 11.0 –

296a WR MIV-SA Ath(10) Cum(5) Qtz(40) Pin(23) Grt(7) Op(10) Chl(5) 17.5 –

635b8b,a WR MIV-WA MgHbl(50) Qtz(30) An(10) Cum(3) Ep(2) Zo(2) Op(3). 16.9 –

635c4b,a WR MIV-TZ MgHblCTs (10) Cum(40) An(13) Qtz(28)

Crd(5) Op(4)

17.4 –

635c1b,a WR MIV-MA/PZ2 Cum(50) Qtz(28) Grt(7) Crd (5) An(5) Op(5). 17.0 –

635hb,a WR MIV-SA Cum(15) Ath(5) Qtz(30) Crd(28) Grt(15) Op(5) Pl(2) 17.8 –

100 SS2 Grt MIV-SA – 16.6 –

100 SS2 Qtz MIV-SA – 19.1 –

288 Grt MIV-SA – 17.0 –

288 Qtz MIV-SA – 19.7 –

Rocks from chloritic alteration zone 2 (CZ2)

127vb,a WR MCI MgHblCTsCPrg(70) Ri(21) Op (7) Pl(2) 9.0 –

217a WR MCI Hbl(65) Chl(20) Pl(7) Op(5) 10.6 –88

Potassified rocks

440b WR MII-TZCPA Hbl(5) Cum(35) Crd(20) Qtz(15) Bt(15) Pl(7) Op(3) 16.2 –

417a WR MII-MACPA CumCAth(35) Qtz(25) Bt(25) Pl(8) Pin(6) Op(1) 15.9 –

Marundites

338b WR Pl marundite Crn(30) Pl(25) Ms(25) Mrg(15) Rt(5) 9.7 –55

Ma-6p Ms Ms schist Ms(98) Op(2) 9.9 –80

Ma-12o Mrg Mrg schist Mrg(98) Op(2) 9.9 –100

Notes: Mineral abbreviations after Kretz (1983); others are as follows: Op, opaque minerals; Pin, pinite; Ri, ripidolite; WR, whole-rock; SS2, syn-to post-S2;

MBI, metabasite/igneous protolith; MBL, metabasite/lapillistone protolith; MBT, metabasite/lapilli-tuff protolit; MCI, metachloritite/igneous protolith; MII,

intermediate metamorphosed rock/igneous protolith; MIV, intermediate metamorphosed rock/volcaniclastic protolith; NA, not affect by hydrothermal

alteration; WA, weakly altered rock; TZ, transitional alteration zone; MA, moderately altered rock; SA, strongly altered rock; PA, potassic alteration; PZ1,

petrographic zone 1; and PZ2, petrographic zone 2.a Microprobe analyses.b Samples from continuous section.c Duplicated values.

A. Perez-Aguilar et al. / Journal of South American Earth Sciences 18 (2005) 305–321312

Fig. 8. d18O values for (1) CZ1 metabasic igneous rocks, (2) CZ1 metahydroclastic rocks, (3) CZ1 intermediate metamorphosed igneous rocks, (4) CZ1

intermediate metamorphosed volcaniclastic rocks, (5) CZ1 continuous outcrop in intermediate metamorphosed volcaniclastic rocks, (6) potassic alteration

overprinting intermediate intermediate metamorphosed igneous rocks from CZ1, and (7) CZ2 basic igneous derived metachloritites.

A. Perez-Aguilar et al. / Journal of South American Earth Sciences 18 (2005) 305–321 313

5.2. Stable isotope mineral compositions

Fig. 9. Calculated oxygen and hydrogen isotopic compositions of fluid in

equilibrium with muscovite and margarite for temperatures between 200

and 350 8C (Suzuoki and Epstein, 1976; Zheng, 1993b). Also shown are the

whole-rock oxygen and hydrogen isotopic compositions of plagioclase

marundite and metachloritite.

5.2.1. Quartz-garnet

The d18O values of the two quartz-garnet pairs from

strongly hydrothermally altered, intermediate metavolcani-

clastic rocks are C19.1 and C16.6‰ (sample 100) and

C19.7 and C17.0‰ (sample 288) (Table 1). Calculated

temperatures of 822 8C (sample 100) and 867 8C (sample

288) were obtained from the oxygen isotopic fractionation

equation between quartz and almandine proposed by Zheng

(1993a) on the basis of the garnet composition obtained by

Perez-Aguilar (2001).

The temperatures, though consistent, are extremely high

and geologically unlikely given the metamorphic peak

temperatures estimated for the Mesoproterozoic (650 8C)

and Neoproterozoic (580 8C) events (Juliani et al., 1997).

Thus, our results could indicate isotope disequilibrium

related to hydrothermal quartz and metamorphic garnet

isotope signatures or reflect the presence of very fine-

grained intergrowths of quartz and garnet in garnet

poikiloblasts (Fig. 7c). Quartz might be a refractory mineral

in relation to stable isotope changes in low water: rock

ratios, typical of those of medium- to high-grade

metamorphic conditions. Assuming a hydrothermal signa-

ture for quartz and temperatures consistent with chloritic

alteration zones in hydrothermal oceanic systems

(200–350 8C, cf. Ohmoto, 1996; Honnorez et al., 1998),

we calculated the d18O values of the hydrothermal fluid in

equilibrium with quartz using the quartz-water fractionation

curves of Clayton et al. (1972); Friedman and O’Neil

(1977); Matsuhisa et al. (1979), and Zheng (1993a). The

calculated d18O values from these different isotopic

fractionation equations are similar (Table 2) and vary

from C5.8‰ (200 8C) to C14.4‰ (350 8C). In addition, a

value of C19.8‰ was calculated using the almandine-water

fractionation curve of Zheng (1993a) at 600 8C for a

probable metamorphic fluid in equilibrium with the

metamorphic garnet (Table 3).

5.2.2. Muscovite and margarite

The d18O and dD values for mineral phases from rocks

associated with marundites are C9.9 and K80‰ (musco-

vite) and C9.9 and K100‰ (margarite), respectively

(Table 1). The oxygen and hydrogen isotope compositions

Table 2

d18O values for water, considering quartz-water oxygen isotopic fraction-

ation at 200–350 8C, calculated according to different authors

d18O fluid (V-SMOW ‰)

Range Zheng

(1993a)

Matsuhisa

et al. (1979)

Friedman and

O’Neil (1977)

Clayton et al.

(1972)

0–1200 8C 250–500 8C 195–573 8C 200–500 8C

Sample 100 288 100 288 100 288 100 288

200 8C 7.5 8.1 7.5 8.1 5.8 6.4 6.9 7.5

250 8C 10.1 10.7 10.2 10.8 8.7 9.3 9.7 10.3

300 8C 12.1 12.7 12.2 12.8 10.9 11.5 11.7 12.3

350 8C 13.5 14.1 13.8 14.4 12.5 13.1 13.3 13.9

A. Perez-Aguilar et al. / Journal of South American Earth Sciences 18 (2005) 305–321314

of the fluid in equilibrium with these minerals were

calculated for a wide range of temperatures (200–600 8C)

using isotopic fractionation equations between muscovite

and water given by of O’Neil et al. (1969); Bottinga and

Javoy (1973); Suzuoki and Epstein (1976), and Zheng

(1993b) (Table 3). In this temperature range, the calculated

d18O values of the hydrothermal fluid in equilibrium

with muscovite, at 200 and 600 8C, vary from C5.4‰ to

C11.1‰, and the calculated dDfluid values range from 0‰

to K73‰ (Table 3). The calculated d18O values of the fluid

in equilibrium with margarite, using the same temperature

range, result in similar d18O and dD values ranging from

K20‰ to K93‰ (Table 3).

6. Discussion

6.1. Whole-rock stable isotope constraints

on the paleohydrothermal system evolution

In a mid-ocean ridge environment, the recharge of

convective cold seawater descends through geothermal

gradients, from high permeability upper volcanics to low

permeability dikes and finally to very low permeability

Table 3

d18O and dD values of fluid phase for 200–600 8C, considering almandine-water,

d18O H2O (‰ V-SMOW)

Almandine-Water Muscovite-Water

T (8C) Zheng (1993a) Zheng (1993b) O’Neil et al.

(1969)

Sample 100 288 Pb-ma-6p Pb-ma-6p

200 16.4 16.8 5.4 (3.2)

250 17.6 18.0 7.2 (5.1)

300 18.4 18.8 8.5 (6.5)

350 18.9 19.3 9.3 (7.7)

400 19.2 19.6 9.9 8.5

450 19.4 19.8 10.3 9.2

500 19.5 19.9 10.6 9.8

550 19.5 19.9 10.8 10.3

575 19.5 19.9 11.0 10.7

600 19.4 19.8 11.1 11.0

Notes: (), data outside of the calibration range.

gabbros. The venting cycle is associated with the presence

of a heat source approximately 2.5 km deep, by which

marine fluids are heated and chemically modified before

their ascending discharge through the rock sequence.

Discharge promotes alteration of wall rocks

and precipitation of massive sulfides (Franklin et al.,

1981; Cathles, 1983; Ohmoto, 1996).

The unaltered metabasites, which have chemical com-

positions typical of N-MORB (Juliani, 1993; Juliani et al.,

2000a), display d18O values between C5.9 and C9.0‰.

The lowest value is similar to those reported for the N- or

E-MORB (d18OZC5.35 to C6.05‰, cf. Ito et al., 1987;

C5.2 to C5.8‰, Eiler et al., 2000; Eiler, 2001), in support

of the preservation of the primary oxygen isotope signature

of this rock despite the regional metamorphic overprint.

Higher values (C7.8 to C9.0‰) could indicate that some

metabasites that appear unaltered were affected by weak

hydrothermal alteration related to submarine discharge

zones, which resulted in small 18O shifts without significant

chemical changes that would crystallize typical

hydrothermal-related metamorphic minerals, such as cum-

mingtonite or cordierite. Juliani (1993) and Perez-Aguilar

(2001) observe a similar situation in whole-rock and trace

element behavior. Alternatively, this weak hydrothermal

alteration may be explained by near-seafloor, low-tempera-

ture isotopic exchange (!150 8C, Muehlenbachs, 1986)

between rocks and seawater (Gregory et al., 1981; Eiler,

2001) before the installation of the hydrothermal systems, as

we discuss subsequently. Typically, seafloor alteration of

basalts results in d18O values of approximately C8 to C9‰

(Muehlenbachs, 1986), but values up to C12.7‰ (Gregory

et al., 1981) and C19.2‰ (Staudigel et al., 1995) also may

be associated with the process.

The oxygen isotope pattern observed in the metabasites

and metamorphosed intermediate igneous rocks from CZ1

represents well-defined trends of increasing d18O values

with the progressive intensity of the alteration process, as is

muscovite-water, and margarite-water pairs

dD H2O (‰ V-SMOW)

Margarite-Water Muscovite-

Water

Margarite-

Water

Bottinga and

Javoy (1973)

Zheng (1993b) Suzuoki and

Epstein (1976)

Suzuoki and

Epstein (1976)

Pb-ma-6p Pb-ma-12o Pb-ma-6p Pb-ma-12o

(4.5) 5.4 0 K20

(6.1) 7.2 K18 K38

(7.2) 8.5 K32 K52

(8.1) 9.3 K42 K62

(8.8) 9.9 K50 K70

9.4 10.3 K57 K77

9.8 10.6 K62 K82

10.2 10.8 K67 K87

10.5 11.0 K70 K90

10.8 11.1 K73 K93

A. Perez-Aguilar et al. / Journal of South American Earth Sciences 18 (2005) 305–321 315

characterized by the mineralogy of these rocks (Table 1). In

banded rocks derived from fine-grained volcaniclastic

protoliths of basaltic andesitic and andesitic composition

from the CZ1, a similar trend, but with a narrow d18O range

(C15.3 to C17.8‰), is identified for almost all samples

(Fig. 8). This narrow range probably is due to the high

porosity, permeability, and relative abundance of volcanic

glass that can occur in volcaniclastic material (Staudigel

et al., 1995), which enables intense fluid–rock interactions

and favors stable isotopic homogenization.

Two features call attention to the d18O values obtained

for hydrothermally altered rocks from the CZ1. The first

is the extremely high whole-rock 18O enrichment (up to

C17.8‰, Table 1) compared with the typical oxygen

isotope signature of unaltered basic and intermediate

igneous rocks (d18OZC5.5 to C11.0‰, cf. Taylor and

Sheppard, 1986); the second is the reverse d18O pattern

compared with wall rocks of most VMSDs.

Relatively high d18O whole-rock values also have been

observed in wall rocks associated with massive sulfide

deposits of Aljustrel in the Carboniferous Iberia pyrite belt

(Spain), the Silurian Blue Hill (Maine, USA), and Kidd

Creek and Mobrun in the Archean Abitibi greenstone belt

(Canada) (Barriga and Kerrich, 1984; Munha et al., 1986;

Beaty et al., 1988; Hoy, 1993). These values have been

interpreted as evidence of high 18O ore-forming fluids and/

or an early low-temperature, near-seafloor alteration stage.

At the Serra do Itaberaba paleohydrothermal system,

extremely high 18O enrichment in rocks could be attributed

to both processes, as we discuss next.

An early low-temperature, essentially fracture

controlled, widespread exchange of seafloor rocks with

marine water at high water: rock ratios could be responsible

for a previous heterogeneous 18O enrichment in the CZ1

rocks, as supported by the relatively high d18O values-C15.5

and C16.9‰, respectively, from samples of unaltered and

weakly altered metamorphosed intermediate volcaniclastic

rocks (291c and 635b8). These samples are extremely18O-enriched compared with the oxygen isotope patterns of

intermediate rocks and can be related to high 18O signatures

of volcanic rocks found peripheral to Kuroko deposits in the

zeolite facies (C16.9G2.7‰, Green et al., 1983; C13 to

C23‰, Ohmoto, 1996). In the upper oceanic crust, alteration

starts immediately after crust formation, is very

heterogeneous, and is controlled by the temperature and

oxidation potential of circulating waters (Bohlke et al., 1981;

Staudigel et al., 1981). Heterogeneous d18O whole-rock

values can be inherited through this process. However, the

strong relationship between alteration intensity and the

oxygen isotope compositions points to an important contri-

bution of 18O-enriched fluids in the hydrothermal system.

The 18O enrichment of seawater fluids can be related to

several processes: (1) mixing with magmatic fluids;

(2) mixing with high 18O connate or metamorphic water

from underlying formations; (3) hydrothermal interaction

with host rocks at high temperatures and a low water: rock

ratio; (4) fluid interaction with 18O-enriched rocks at high

temperatures and a low water: rock ratio; (5) seawater

convection through footwall-high 18O sediments; (6) multi-

pass, semiclosed-system seawater convection; (7) shale

ultrafiltration; (8) seawater evaporation in a closed basin;

and (9) hydrothermal boiling (Munha et al., 1986). The

tectonic evolution of the Serra do Itaberaba Group from an

oceanic (N-MORB) to a backarc basin environment

suggests that the plausible mechanisms for 18O-enriched

fluids are as follows: high temperature and low water: rock

interaction of fluids with previously 18O-enriched rocks,

evaporation in a closed basin, hydrothermal boiling, shale

ultrafiltration, convection through footwall-high 18O

sediments, and magmatic fluid contribution.

In the CZ1 rocks, relatively well-defined trends of

increasing d18O values with the progressive intensity of the

alteration process are observed. The d18O halos are opposite

those of most Archean and Phanerozic VMSDs, including

Kuroko-type deposits, in which 18O values decrease toward

the mineralized zone and increase outward from this zone

(Barrett and MacLean, 1994; Ohmoto, 1996; Vasquez et al.,

1998; Gemmell et al., 1998; Cartwright, 1999; Shikazono,

2003). However, at the Kidd Creek VMSD, a reverse d18O

pattern, similar to that of the Serra do Itaberaba Group, is

observed at both a decimeter scale beneath the mineraliz-

ation zone and a regional scale with higher whole-rock 18O

values near the mineralized zone. This feature implies at

least a two-stage hydrothermal evolution (Beaty et al.,

1988).

In the CZ1 igneous and volcaniclastic rocks, continuous

alteration from weakly to strongly altered rocks is observed

at the meter to decimeter scale (Figs. 5 and 6). In addition,

protolith textures are relatively preserved in weakly to

moderately altered rocks but completely destroyed in

strongly altered rocks, substituted for by characteristic

arrays of radiate magnesium amphiboles and garnet and/or

cordierite poikiloblasts. Variations in fluid temperatures or

fluid compositions within a single hydrothermal event

cannot explain these features. The textures suggest that

weak to moderate alteration took place in low water: rock

ratio conditions, whereas strong alteration occurred in

water-dominant conditions, mainly in highly permeable

rocks that served as preferred channelways for hydrothermal

fluids. Thus, the reverse d18O pattern in the CZ1 rocks could

be a consequence of the variable original permeability of the

protoliths, which would favor different water: rock ratios in

the volcanosedimentary layers or igneous bodies. As a

consequence, hydrothermal activity promoted variable

degrees of 18O enrichment in the rocks relative to the

assumed primary d18O values of basic and intermediate

rocks (C5.5 to C11.0‰, Taylor and Sheppard, 1986).

Isotope rock d18O uniformity requires high permeability

(Barriga and Kerrich, 1984), which enables a significant

flow of water through the rocks (high water: rock ratios) and

a pervasive alteration style. In the CZ1 rocks, these

conditions prevailed only during the formation of

A. Perez-Aguilar et al. / Journal of South American Earth Sciences 18 (2005) 305–321316

the strongly altered rocks, to which should correspond to

similar d18O values. However, lower d18O values are

observed in a strongly altered intermediate metavolcani-

clastic rock (sample 301a, d18OZC11.6‰,) and a strongly

altered metabasite (sample 127b, d18OZ11.8‰). These

lower values might correspond to a different subordinate

hydrothermal pulse alteration at a high water: rock ratio but

higher temperature or lower d18O fluid composition.

Rocks from the CZ2, which are restricted to feeder zones,

show lower d18O values (C9.0 to C10.6‰) in comparison

with most of the hydrothermally altered rocks from the CZ1

(Fig. 8). A similar pattern also is observed in most VMSDs

(Green et al., 1983; Ohmoto, 1996; Gemmell et al., 1998;

Cartwright, 1999; Shikazono, 2003). Alteration of CZ2

rocks must have been produced by higher temperature fluids

compared with those responsible for the alteration of rocks

from the CZ1, which caused oxygen isotopic shifts toward

lower values.

The d18O and dD whole-rock data of a metachloritite

from the CZ2 are C10.6 and K88‰ (sample 217a) and of a

plagioclase marundite C9.7 and K55‰, respectively

(Table 1). These stable isotope compositions plot in and

near the field of primary igneous rocks (Fig. 9). The

hydrogen isotope composition of the metachloritite is

similar to MORB values (K80G5‰, cf. Craig and Lupton,

1976; Kyser and O’Neil, 1984; Alt et al., 1996). The

negative dD values suggest the contribution of magmatic

water (dDZK40 to K80‰; cf. Sheppard et al., 1979; Rye,

1993) to the hydrothermal alteration system responsible for

the generation of chloritic (CZ2), argillic, and advanced

argillic alteration zones.

The oxygen isotope signature of plagioclase marundite is

relatively 18O depleted in relation to the CZ1 rocks but

similar to that of the metachloritite from the CZ2 (Table 1).

This finding is consistent with petrographic evidence of

premetamorphic replacement of volcaniclastic fragments in

basic tuffs by argillic and advanced argillic alteration clay

products that, after metamorphism, led to the formation of

marundites (Juliani, 1993).

High-alumina clays derived from the hydrothermal

alteration of volcanic or intrusive rocks are present in

Kuroko-type mineralized zones from Japan and Korea

(Schmidt, 1985; Shikazono, 2003). These clay rocks are

associated with intermediate or felsic intrusive bodies,

which constitute the heat engines that drive the hydrother-

mal circulation (Schmidt, 1985) and are formed in high-

(300–400 8C) and low- (100–300 8C) temperature argillic

and advanced argillic alteration zones. At low temperatures,

the mineral assemblage of the altered rocks is

typically kaoliniteCpyrophylliteCaluniteCquartzGdia-

sporeGpyrite (Meyer and Hemley, 1967; Hemley et al.,

1980; Silberman and Berger, 1985), which could generate

marundites with medium-grade metamorphism (Juliani,

1993; Juliani et al., 1994).

Thus, the Serra do Itaberaba marundites may represent

the metamorphic product of low-temperature argillic and

advanced argillic hydrothermal alteration generated by acid

and sulfate-rich fluids circulating near the ocean floor and

associated with andesitic to rhyodacitic intrusions. Musco-

vite and margarite schists, typically without quartz,

commonly are associated with marundites (Juliani, 1993;

Juliani et al., 1994; Perez-Aguilar, 2001). These rocks could

be products of changing KC, Al3C, and Ca2C activities in

the hydrothermal fluids, which would favor sericitic

alteration zones similar to those that envelope the chloritic

alteration zone in Kuroko-type deposits (Green et al., 1983;

Ohmoto, 1996; Shikazono, 2003).

The d18O values from chloritized, intermediate, moder-

ately altered igneous rocks from the transitional zone,

overprinted by potassic alteration (C15.9 and C16.2‰),

suggest that K metasomatism (Fig. 7b) did not substantially

modify the oxygen isotope signature related to CZ1

alteration.

6.2. Preservation of the hydrothermal oxygen

isotope signatures

The well-defined oxygen isotope trends indicate that the

original hydrothermal system isotope signatures were

preserved and perhaps that the different superimposed

metamorphic events did not promote significant homogen-

ization of the preexisting isotope patterns. Aggarwal and

Longstaffe (1987) interpret differences in the oxygen

isotope compositions of altered host rocks of metamor-

phosed massive sulfide deposits in the Flin Flon-Snow Lake

belt (Canada) as produced before metamorphism, during

hydrothermal alteration related to ore deposition. Preser-

vation of hydrothermal oxygen isotope signatures in

Precambrian massive sulfide deposits was observed by

Beaty et al. (1988); Beaty and Taylor (1982), and Araujo

et al. (1996), in which cases dehydration reactions did not

significantly affect d18O values.

6.3. Oxygen and hydrogen isotope fluid compositions

Quartz included in syn- to post-S2 garnet poikiloblasts

from strongly hydrothermally altered intermediate volcani-

clastic rocks is enriched in 18O (samples 100 and 288,

d18OZC19.1 and C19.7‰, respectively) relative to most

igneous quartz (C8 to C12‰; Taylor, 1968). This strong18O enrichment of quartz was acquired through exchange

with an 18O-enriched pervasive fluid, as we discussed

previously.

The d18O values for the fluid in equilibrium with quartz

from the CZ1 (C5.8 to C14.4‰; Table 2) were calculated

for 200–350 8C, consistent with temperatures present in

chloritic alteration zones of VHMS deposits (Ohmoto,

1996; Shikazono, 2003). At these temperatures, the d18O

fluid values, which are higher than those of normal ocean

waters (Muehlenbachs, 1986).

The highest d18O fluid calculated from a garnet-water

pair (Table 3, Fig. 10) may reflect a predominant

Fig. 10. Calculated fluid oxygen isotopic composition at 300 8C, in

equilibrium with quartz, garnet, muscovite, and margarite.

A. Perez-Aguilar et al. / Journal of South American Earth Sciences 18 (2005) 305–321 317

metamorphic signature. However, the calculated isotope

composition also could be due to quartz-garnet intergrowth,

which might increase the measured d18Ogarnet values and,

consequently, the calculated d18O of the fluid. However, this

effect would be rather small (less than 1‰).

The calculated oxygen and hydrogen fluid composition

(Table 3) in equilibrium with muscovite (d18OZC5.4 to

C8.5‰; dDZ0 to K32‰) and margarite (d18OZC5.4 to

C8.5‰; dDZK20 to K52‰) at temperatures consistent

with those typical of the premetamorphic low temperature

argillic and advanced argillic alteration (200–300 8C) are

higher than the d18O seawater values (Fig. 9). The negative

dD values obtained for the fluid phase in equilibrium with

margarite and muscovite suggest mixing that involved

mostly seawater with a smaller contribution of magmatic

water (d18OZC5.5 to C9.5‰; dDZK40 to K80‰;

Sheppard et al., 1979). This magmatic water contribution

could be related to the shallow intrusion of the andesitic and

rhyodacitic bodies in the backarc basin environment.

Evaporation in a closed basin, seawater convection through

footwall-high 18O sediments, shale ultrafiltration, and

seawater boiling would increase the dD and d18O values

of the residual fluids. However, the calculated negative dD

values of the fluids imply that these mechanisms cannot

explain the observed high 18O character of hydrothermal

fluids related to argillic and advanced argillic alteration.

The oxygen isotope fluid composition in equilibrium

with quartz, garnet, muscovite, and margarite at 300 8C

indicates that the fluids related to CZ1 alteration were more18O enriched than those related to the argillic and advanced

argillic alteration zones (Fig. 10).

We suggest that highly evolved hot seawater fluid,

inherited from high temperature and low water: rock ratio

interaction with 18O-enriched rocks, generated the CZ1

rocks, whereas for the CZ2 rocks and marundites, the 18O

fluid enrichment represents a major contribution of evolved

seawater-derived fluids with a subordinate magmatic fluid

component.

There are several examples of 18O-enriched fluids

associated with VMSDs, such as the Archean Kidd Creek

deposit, Canada (d18OZC6.0 to C9.0‰; Beaty et al.,

1988); the Ordovician Heath Steele B zone deposit, New

Brunswick, Canada (d18OZC5.0 to C7.0‰; Lentz et al.,

1997); the Silurian Blue Hill deposits, USA (d18OZC5.0 to

C6.6‰; Munha et al., 1986); the Triassic Gacun deposit,

Sichuan, China (d18OZC5.5 to C8.5‰; Zengquian et al.,

2001); and the Cretaceous Raul mine, Peru (d18OZC9.1 to

C12.6‰; Ripley and Ohmoto, 1979).

Hoy (1993) points out that in the Noranda district, the

d18O values of the altered rocks and the economic tonnage

of each deposit increase upward through the volcanic

stratigraphy from low d18O values at the Corbet (K2.2 to C4.8‰) and Ansil (K0.8 to C5.0‰) deposits to intermediate

d18O values at the Amulet (C3.6 to C6.7‰) and Norbec

(C3.6 to C10.5‰) deposits and then to high d18O values at

the Horne (C4.2 to C11.6‰) and Mobrun (C6.0 to C13.8‰) deposits. These increasing d18O values for altered

rocks correspond to a progressive 18O enrichment of the

hydrothermal fluids from K2.0G2‰ at Corbet to C3.0G1.5‰ at Horne. Hoy (1993) suggests that hydrothermal

discharge duration is correlated with the size of the sulfide

ore bodies and apparently is a primary control on the isotope

composition of the rocks and mineralizing fluids.

Beaty et al. (1988) suggest that high 18O ore-forming

fluids are characteristic of exceptionally large base metal

deposits. This hypothesis has been refuted by Munha et al.

(1986) on the basis of the presence of heavy d18Ofluid values

recorded at the relatively small Blue Hill deposits.

In several “super giant” massive sulfide deposits

(Ohmoto, 1996)—such as the Archean Kidd Creek and

Horne deposits in the Abitibi belt (Canada); the Carbon-

iferous Rio Tinto, San Guillermo, Filon, La Zarza, and

Ajustrel deposits in the Iberian pyrite belt (Europe); and the

Proterozic Crandon deposit in Wisconsin (USA) (Munha

et al., 1986; Beaty et al., 1988; Hoy, 1993)—one or more of

the following features are present: an early low-temperature

seafloor alteration stage, high d18O values of wall rock, high18O ore-forming fluids, and a lack of oxygen isotope

homogenization in wall rocks. These characteristics point to

long-lived hydrothermal systems, which may generate large

base metal sulfide deposits. Therefore, the stable isotope

signatures of the multiple hydrothermal events that occurred

in the Serra do Itaberaba Group suggest an effective, long-

lived hydrothermal system that could generate massive

sulfide ore bodies in addition to the known, extensive gold

mineralization (Garda et al., 2002).

7. Conclusions

A multistage hydrothermal history is recognized in the

Mesoproterozoic paleohydrothermal systems of the Serra do

Itaberaba Group. An early hydrothermal activity stage,

related to nonsteady-state convective cooling of the oceanic

A. Perez-Aguilar et al. / Journal of South American Earth Sciences 18 (2005) 305–321318

crust, was responsible for the heterogeneous d18O enrich-

ment of seafloor rocks (up to C15.5‰). During backarc

basin evolution, the emplacement of shallow andesitic and

rhyodacitic intrusions was responsible for the development

of a long-lived hydrothermal system, which resulted in

large, external, premetamorphic chloritic (CZ1 rocks)

alteration halos and intensely leached rocks in the feeder

zones, characterized by restricted chloritic (CZ2 rocks),

argillic, and advanced argillic (marundites) alteration,

similar to Kuroko-type deposits. This hydrothermal system

was related to the discharge of fluids previously heated in

the deeper parts of the system, where they underwent

isotopic exchange with 18O-enriched volcanic rocks at high

temperatures and low water: rock ratios.

The metamorphic products of these hydrothermalized

zones are typical rocks composed of cummingtonite,

anthophyllite, gedrite, cordierite, garnet, Mg-chlorite,

staurolite, ilmenite, rutile, Ca-rich plagioclase, and quartz

(CZ1 rocks); rocks composed of magnesiohornblendeCtschermakiteGMg-chlorite or cummingtoniteGMg-

chloriteGgarnetGplagioclase (CZ2 rocks); and corundum-,

margarite- and rutile-bearing rocks (marundites).

Systematic stable isotope variations, represented by

well-defined trends of increasing values of d18O with

progressive alteration process intensity, provide evidence

that the preexisting isotope signatures of the hydrothermal

systems were preserved, despite two medium-grade

metamorphic-deformational events that affected the rocks

during the Meso- and Neoproterozoic. These trends are the

consequence of hydrothermal activity controlled mainly by

rock permeability and water: rock ratios.

Integrated geological, petrological, and oxygen and

hydrogen isotope evidence points to the participation of

highly evolved hot seawater related to the high-temperature

and low water: rock ratio interaction with 18O-enriched

rocks for the genesis of the CZ1 rocks. For the CZ2 rocks

and marundites, in contrast, the 18O fluid enrichment

represents a major contribution of evolved seawater-derived

fluids, with a subordinate magmatic fluid component

derived from the andesitic and rhyodacitic intrusions.

The extremely high d18O anomalies given by rocks from

the CZ1 and related hydrothermal fluids were achieved

through a long-lived hydrothermal system, which suggests

the possibility that base metal massive sulfide ore bodies

associated with this hydrothermal activity exist, in addition

to extensive gold mineralizations (Garda et al., 2002). The

relatively lower d18O values related to the metachloritites

(CZ2) represent alteration by a higher temperature fluid,

which may indicate proximity of the feeder zones related to

the possible ore zones.

In this context, activity related to black smokers (Fouquet

et al., 1991) may correspond to the metalliferous metase-

dimentary rocks, especially sulfide-rich ones that occur

along the interface of the Morro da Pedra Preta and the

Nhangucu Formations and are characterized by gold

mineralization and soil copper and zinc anomalies (Juliani

et al., 1986).

Hydrothermal alteration halos are usually larger than the

ore bodies. The preservation of premetamorphic stable

isotope signatures in the hydrothermalized Serra do

Itaberaba Group rocks shows that, despite medium-grade

metamorphic events, stable isotopes can be used for mineral

exploration, especially for nonexposed ore bodies. The

identification of a large premetamorphic hydrothermal halo,

similar to those of Kuroko-type base metal mineralization,

expands the potential for base metal deposits in the Serra do

Itaberaba Group and the volcanosedimentary sequences

from the Ribeira fold belt.

Acknowledgements

The authors thank the Fundacao de Amparo a Pesquisa

do Estado de Sao Paulo (grants 93/4350-0 and 98/15170-7)

and the Conselho Nacional de Desenvolimento Cientıfico e

Tecnologico for research grant 400490-94-3 and for the

Masters and Ph.D. scholarships granted to Annabel Perez-

Aguilar. The authors also are grateful to reviewers Sylvia

Maria Araujo, Reinhardt Fuck, and Philip Piccoli, who

significantly improved this article.

References

Aggarwal, P.K., Longstaffe, F.J., 1987. Oxygen-isotope geochemistry of

metamorphosed massive sulfide deposits in the Flin Flon-Snow Lake

Belt, Manitoba. Contributions to Mineralogy and Petrology 96,

314–325.

Almeida, F.F.M., de Amaral, G., Cordani, U.G., Kawashita, K., 1973. In:

Nairn, A.E.M., Stehli, F.G. (Eds.), The Precambrian evolution of the

South American cratonic margin south of Amazon River The Ocean

Basin and Margins, 1. Plenum Press, New York, pp. 411–446.

Almeida, F.F.M., Hasui, Y., Poncano, W.L., Dantas, A.S.L.,

Carneiro, C.D.R., de Melo, M.S., Bistrichi, C.A., 1981. Mapa geologico

do Estado de Sao Paulo, escala 1:500.000-Nota Explicativa. Sao Paulo,

escala 1:500.000-Nota Explicativa. Sao Paulo, p. 6.

Alt, J.C., Laverne, C., Vanko, D., Tartarotti, P., Teagle, D.A.H., Bach, W.,

Zuleger, E., Erzinger, J., Honnorez, J., Pezard, P.A., Becker, K.,

Salisbury, M.H., 1996. In: Alt, J.C., Kinoshita, H., Stokking, L.B.,

Michael, P.J. (Eds.), Hydrothermal alteration of a section of upper

oceanic crust in the eastern equatorial Pacific: a synthesis of results

from Site 504 (DSDP Legs 69, 70, and 83, and ODP Legs 111, 137, 140,

and 148) ODP proceedings, science results, 148, pp. 417–434.

Araujo, S.M., Scott, S.D., Longstaffe, F.J., 1996. Oxygen isotope

composition of alteration zones of highly metamorphosed volcanogenic

massive sulfide deposits: Geco Canada, and Palmeiropolis, Brasil.

Economic Geology 91, 697–712.

Barrett, T.J., MacLean, W.H., 1994. Mass changes in hydrothermal

alteration zones associated with VHMS deposits in Noranda area.

Exploration and Mining Geology 3, 131–160.

Barriga, F.J.A.S., Kerrich, R., 1984. Extreme 18O-enriched volcanics and18O-evolved marine water Aljustrel, Iberian pyrite belt: Transition from

high to low Rayleigh number convective regimes. Geochimica et

Cosmochimica Acta 48, 1021–1031.

A. Perez-Aguilar et al. / Journal of South American Earth Sciences 18 (2005) 305–321 319

Beaty, D.W., Taylor Jr.., H.P., 1982. Some petrologic and oxygen isotopic

relationships in the Amulet mine Noranda, Quebec, and their bearing on

the origin of Archean massive sulfide deposits. Economic Geology 77,

95–108.

Beaty, D.W., Taylor, Jr., H, P., Coad, P.R., 1988. An oxygen isotope study

of the Kidd Creek Ontario, volcanogenic massive sulfide deposit:

evidence for high 18O ore fluid. Economic Geology 83, 1–17.

Bohlke, J.K., Honnorez, J., Honnorez-Guersyein, B.-M., Muehlenbachs, K.,

Petersen, N., 1981. Heterogeneous alteration of the upper oceanic crust:

Correlation of rock chemistry, magmatic properties and O isotope ratios

with alteration patterns in basalts from site 396B, DSDP. Journal of

Geophysical Research 86, 7935–7950.

Bottinga, Y., Javoy, M., 1973. Comments on oxygen isotope geothermo-

metry. Earth and Planetary Science Letters 20, 250–265.

Cartwright, I., 1999. Regional isotope zonation at Broken Hill New South

Wales, Australia: large scale fluid flow and implications for Pb–Zn–Ag

mineralizations. Economic Geology 94, 357–374.

Cathles, L.M., 1983. An analysis of the hydrothermal system responsible

for massive sulfide deposition in the Hokuroku basin of Japan.

Economic Geology Monograph 5, 439–487.

Cathles, L.M., 1993. Oxygen isotope alteration in the Noranda mining

district Abitibi greenstone belt, Quebec. Economic Geology 88, 1483–

1511.

Clayton, R.N., O’Neil, J.R., Mayeda, T.K., 1972. Oxygen isotope exchange

between quartz and water. Journal of Geophysical Research 77, 3057–

3067.

Craig, H., Lupton, J.E., 1976. Primordial neon, helium, and hydrogen in

oceanic basalts. Earth and Planetary Science Letters 31, 369–385.

Criss, R.E., Taylor Jr., H.P., 1983. An 18O/16O and D/H study of tertiary

hydrothermal systems in the southern half of the Idaho batholith.

Geological Society of America Bulletin 94, 640–663.

Dobbe, R.T.M., 1994. Geochemistry of cordierite-anthophyllite rocks

Tunaberg, Bergslagen, Sweden. Economic Geology 89, 919–930.

Donnelly, T., Waldron, S., Tait, A., Dougans, J., Bearhop, S., 2001.

Hydrogen isotope analyses of natural abundance and deuterium-

enriched waters by reduction over chromium on-line to a dynamic

dual inlet isotope-ratio mass spectrometer. Rapid Communication in

Mass Spectrometry 41, 184–190.

Eiler, J.M., 2001. In: Valley, J.W., Cole, D.R. (Eds.), Oxygen isotope

variations of basaltic lavas and upper mantle rocks Stable Isotope

Geochemistry. Mineralogical Society of America, Reviews in Miner-

alogy and Geochemistry, 43, pp. 319–364.

Eiler, J.M., Schiano, P., Kitchen, N., Stolper, E.M., 2000. Oxygen isotope

evidence for recycled crust in the sources of mid ocean ridge basalts.

Nature 403, 530–534.

Elliott-Meadows, S.R., Appleyard, E.C., 1991. The alteration geochemistry

and petrology of the Lau Cu-Zn deposit Lynn Lake area, Manitoba,

Canada. Economic Geology 86, 486–505.

Fallick, A.E., Macaulay, C.I., Haszeldine, R.S., 1993. Implications of

linearly correlated oxygen and hydrogen isotopic compositions for

kaolinite and illite in the Magnus Sandstone North Sea. Clays and Clay

Minerals 41, 184–190.

Fouquet, Y., Von Sackelberg, U., Charlou, J.L., Donval, J.P., Erzinger, J.,

Foucher, J.P., Herzig, P., Muhe, R., Soakai, S., Wiedicke, M.,

Whitechurch, H., 1991. Hydrothermal activity and metalogenesis in

the Lau back-arc basin. Nature 349, 778–781.

Franklin, J.M., 1993. In: Kirkham, R.V., Sinclair, W.D., Thorpe, R.I.,

Duke, J.M. (Eds.), Volcanic-associated massive sulphide deposits

Mineral Deposit Modeling. Geological Association of Canada Special

Paper, 40, pp. 315–334.

Franklin, J.M., Lydon, J.W., Sangster, D.F., 1981. Volcanic-associated

massive sulfide deposits. Economic Geology 75, 485–627.

Friedman, I., O’Neil, J.R., 1977. Compilation of stable isotope fractionation

factors of geochemical interest, Geological Survey Professional Paper

440-kk 1977.

Garda, G.M., Beljavskis, P., Juliani, C., Boyce, A.J., 2002. Sulfur stable

isotope signatures of the Morro da Pedra Preta Formation Serra do

Itaberaba Group, Sao Paulo State, Brazil. Geochimica Brasiliensis 16,

79–97.

Gemmell, J.B., Large, R.R., Zaw, K., 1998. Palaeozoic volcanic-hosted

massive sulphide deposits. Journal of Australian Geology and

Geophysics 17, 129–137.

Green, G.E., Ohmoto, H., Date, J., Takahashi, T., 1983. Whole-rock oxygen

isotope distribution in the Fukazawa-Kosaka area, Hokuroku district,

Japan and its potential application to mineral exploration. Economic

Geology Monograph 5, 395–411.

Gregory, R.T., Taylor Jr., H.P., 1981. An oxygen isotope profile in a section

of Cretaceous oceanic crust Samail ophiolite, Oman: Evidence for (18O

buffering of the oceans by deep (O5 km) seawater-hydrothermal

circulation at mid-ocean ridges. Journal of Geophysical Research 86,

2737–2755.

Hackspacher, P., Dantas, E.L., Godoy, A.M., Oliveira, M.A.F., de

Fetter, A., Van Schmus, W.R., 1999. Considerations about the evolution

of the Ribeira Belt in the Sao Paulo State—Brazil, from U/pb

geochronology in metavolcanic rocks of the Sao Roque Group, South

American Symposium on Isotope Geology 2, Vila Carlos Paz.

Argentina. Anais, pp. 192–195.

Hall, A.L., 1920. Corundum in the northern and eastern Transvaal.

Geological Survey of South Africa Memoir 1920;, 15.

Hemley, J.J., Montoya, J.W., Marinenko, J.W., Luce, R.W., 1980.

Equilibria in the system Al2O3–SiO2–H2O and some general appli-

cations for alteration/mineralization processes. Economic Geology 75,

210–228.

Honnorez, J.J., Alt, J.C., Humphris, S.E., 1998. In: Herzig, P.M.,

Humphris, S.E., Miller, D.J., Zierenberg, R.A. (Eds.), Vivisection and

autopsy of active and fossil hydrothermal alterations of basalt beneath

and within the TAG hydrothermal mound Proceedings of Ocean

Drilling Program, Scientific Results, 158, pp. 231–254.

Hoy, L.D., 1993. Regional evolution of the hydrothermal fluids in the

Noranda district Quebec: Evidence from (18O values from volcanogenic

massive sulfide deposits. Economic Geology 88, 1526–1541.

Ishihara, S., Kanehira, K., Sasaki, A., Sato, T., Shimazaki, Y., 1974.

Geology of the Kuroko deposits. The Society of Mining Geologists of

Japan, Mining Geology Special Issue 1974;, 6.

Ito, E., White, W.M., Gopel, C., 1987. The O, Sr, Nd and Pb isotope

geochemistry of MORB. Chemical Geology 62, 157–176.

James, R.S., Grieve, R.A.F., Pauk, L., 1978. The petrology of cordierite-

anthophyllite gneisses and associated mafic and pelitic gneisses at

Manitouwadge Ontario. American Journal of Science 278, 41–63.

Juliani, C, 1993. Geologia, petrogenese e aspectos metalogeneticos dos

grupos Serra do Itaberaba e Sao Roque na regiao das serras do Itaberaba

e da Pedra Branca, NE da cidade de Sao Paulo. Unpublished PhD

Thesis, Instituto de Geociencias, Universidade de Sao Paulo, Sao Paulo,

Brazil.

Juliani, C., Beljavskis, P., 1995. Revisao da litoestratigrafia da faixa Sao

Roque/Serra do Itaberaba (SP). Revista do Instituto Geologico 16, 33–

58.

Juliani, C., Beljavskis, P., Schorscher, H.D., 1986. Petrogenese do

vulcanismo e aspectos metalogeneticos associados: Grupo Serra de

Itaberaba na regiao do Sao Roque-SP 34th Congresso Brasileiro de

Geologia, Goiania, Anais, 2, pp. 730–743.

Juliani, C., Schorscher, H.D., Perez-Aguilar, A., Beljavskis, P., 1992.

Cordierita-granada-cummingtonita anfibolitos no Grupo Serra do

Itaberaba (SP): Evidencia de alteracoes hidrotermais-metassomaticas

pre-metamorficas. 2 Jornadas Cientıficas, Boletim do Insituto de

Geociencias da Universidade de Sao Paulo Special Publication 12,

59–61.

Juliani, C., Schorscher, H.D., Perez-Aguilar, A., 1994. Corundum-

margarite schists (“marundites”) in the Precambrian Serra do Itaberaba

Group Sao Paulo, Brazil: geological relationships and petrogenesis.

Anais da Academia Brasileira de Ciencias 66, 498.

A. Perez-Aguilar et al. / Journal of South American Earth Sciences 18 (2005) 305–321320

Juliani, C., Perez-Aguilar, A., Martin, M.A.B., 1997. Geotermobarometria

e evolucao metamorfica P-T-d do Grupo Serra do Itaberaba (SP). Anais

da Academia Brasileira de Ciencias 69, 441–442.

Juliani, C., Hackspacher, P.C., Dantas, E.L., Fetter, A.H., 2000a. The

mesoproterozoic volcano-sedimentary Serra do Itaberaba Group of the

Central Ribeira Belt Sao Paulo, Brazil: implications for the age of

overlying Sao Roque Group. Revista Brasileira de Geociencias 30, 82–

86.

Juliani, C., Hackspacher, P.C., Dantas, E.L., 2000b. Serra do Itaberaba

Group: a Mesoproterozoic volcano-sedimentary unit at Sao Paulo State,

Brazil. 31st International Geological Congress, [CD-ROM].

Kretz, R., 1983. Symbols for rock-forming minerals. American Miner-

alogist 68, 277–279.

Kyser, K.T., O’Neil, J.R., 1984. Hydrogen isotope systematics of

submarine basalts. Geochimica et Cosmochimica Acta 48, 2123–2133.

Lentz, D.R., Hall, D.C., Hoy, L.D., 1997. Chemoestratigraphic, alteration,

and oxygen isotopic trends in a profile through the stratigraphic

sequence hosting the Heath Steele B zone massive sufide deposit, New

Brunswick. The Canadian Mineralogist 35, 841–874.

MacGeehan, P.J., MacLean, W.H., 1980. Tholeiitic basalt-rhyolite

magmatism and massive sulphide deposits at Matagami Quebec.

Nature 283, 153–157.

Matsuhisa, Y., Goldsmith, J.R., Clayton, R.N., 1979. Oxygen isotopic

fractionation in the system quartz-albite-anorthite-water. Geochimica et

Cosmochimica Acta 43, 1131–1140.

Meyer, C., Hemley, J.J., 1967. Wall rock alteration, In: Barnes, H.L. (Ed.),

Geochemistry of Hydrothermal Ore Deposits, first ed. Holt Rinehart and

Wintson, New York, pp. 166–235.

Muehlenbachs, K., 1986. In: Valley, J.W., Taylor Jr., H.P., O’Neil, J.R.

(Eds.), Alteration of oceanic crust and the 18O history of seawater.

Stable isotopes in high temperature geological processes Miner-

alogical Society of America, Reviews in Mineralogy, 16, pp. 425–

444.

Munha, J., Barriga, F.J.A.S., Kerrich, R., 1986. High 18O ore-forming fluids

in volcanic-hosted base metal massive sulfide deposits: geologic,18O/16O, and D/H evidence from the Iberian pyrite belt; Crandon,

Wisconsin; Blue Hill, Maine. Economic Geology 81, 530–552.

Ohmoto, H., 1996. Formation of volcanogenic massive sulfide deposits: the

Kuroko perspective. Ore Reviews 10, 135–177.

O’Neil, J.R., Taylor, J.r., H, P., 1969. Oxygen isotope equilibrium between

muscovite and water. Journal of Geophysical Research 74, 6012–6022.

Pan, Y., Fleet, M.E., 1995. Geochemistry and origin of cordierite-

orthoamphibole gneiss and associated rocks at an Archaean volcano-

genic massive sulphide camp: Manitouwadge, Ontario, Canada.

Precambrian Research 74, 73–89.

Perez-Aguilar, A. Geologia, petrografia e genese dos granada-cordierita-

cummingtonita/antofilita anfibolitos e rochas associadas do Grupo Serra

do Itaberaba, SP. Unpublished Master Thesis, Instituto de Geociencias,

Universidade de Sao Paulo, Sao Paulo, Brazil; 1996.

Perez-Aguilar, A. Petrologia e litoquımica de rochas de paleossistemas

hidrotermais oceanicos mesoproterozoicos da sequencia metavulca-

nossedimentar do Grupo Serra do Itaberaba, SP. Unpublished Ph.D.

Thesis, Instituto de Geociencias, Universidade de Sao Paulo, Sao Paulo,

Brazil; 2001.

Perez-Aguilar, A., Juliani, C., Martin, M.A.B., 2000. Mesoproterozoic

paleo-hydrothermal system in the Morro da Pedra Preta Formation

Serra do Itaberaba Group, Sao Paulo State, Brazil. Revista Brasileira de

Geociencias, 30, pp. 413–416.

Perez-Aguilar, A., Juliani, C., Monteiro, L.V.S., Bettencourt, J.S.,

McReath, I., 2002a. Litoquımica e isotopos estaveis de rochas de

paleossistemas hidrotermais oceanicos Mesoproterozoicos do Grupo

Serra do Itaberaba, SP. Congresso Brasileiro de Geologia, Joao Pessoa,

Anais 41, 485.

Perez-Aguilar, A., Juliani, C., Monteiro, L.V.S., Bettencourt, J.S., 2002b.

Isotopos estaveis de rochas de paleossistemas hidrotermais oceanicos

mesoproterozoicos do Grupo Serra do Itaberaba, SP. Simposio Sobre

Vulcanismo e Ambientes Associados, Belem, Boletim de Resumos

2, 44.

Ripley, E.M., Ohmoto, H., 1979. Oxygen and hydrogen isotopic studies of

ore deposition and metamorphism at the Raul mine Peru. Geochimica et

Cosmochimica Acta 43, 1633–1643.

Riverin, G., Hodgson, C.J., 1980. Wall-rock alteration at the

Millenbach Cu–Zn mine, Noranda, Quebec. Economic Geology

75, 424–444.

Roberts, M.D., Oliver, N.H.S., Fairclouch, M.C., Holita, P.S., Lahtinen, R.,

2003. Geochemical and oxygen isotope signature of sea-floor alteration

associated with a polydeformed and highly metamorphosed massive

sulfide deposit, Ruostesuo, Central Finland. Economic Geology 98,

535–556.

Rye, R.O., 1993. The evolution of magmatic fluids in the epithermal

environment-the stable isotope perspective. Economic Geology 88,

733–753.

Sangster, D.F., 1972. Precambrian volcanogenic massive sulphide

deposits in Canada: a review. Geological Survey of Canada, Paper

1972;, 72.

Schermerhorn, L.J.G., 1978. Epigenetic magnesiam metasomatism or

syngenetic chloritite metamorphism al Falun and Orijarvi.

Institution of Mining and Metallurgy Transactions, Section B

87, 162–167.

Schmidt, R.G., 1985. High-alumina hydrothermal systems in volcanic rocks

and their significance to mineral prospecting in the Carolina slate belt.

United States Geological Survey Bulletin 1985;, 1562.

Sharp, Z.D., 1990. A laser-based microanalytical method for the in situ

determination of oxygen isotopes ratios in silicates and oxides.

Geochimica et Cosmochimica Acta 54, 1353–1357.

Sheppard, S.M.F., Nielsen, R.L., Taylor, J.r., H, P., 1979. Oxygen and

hydrogen isotope ratios of clay minerals from porphyry copper deposits.

Economic Geology 64, 755–777.

Shikazono, N., 2003. Geochemical and tectonic evolution of arc-backarc

hydrothermal systems: Implication for the origin of Kuroko and

epithermal vein-type mineralizations and the global geochemical cycle.

Developments in Geochemistry 2003;, 8.

Silberman, M.L., Berger, B.R., 1985. In: Berger, B.R., Bethke, P.M.

(Eds.), Relationship of trace-element patterns to alteration and

morphology in epithermal precious-metal deposits Geology and

Geochemistry of Epithermal Systems. Reviews in Economic

Geology, 2, pp. 203–232.

Spear, F.S., 1982. Phase equilibria of amphibolites from the post pond

volcanics, Mt. Cube Quadrangle, Vermont. Journal of Petrology 23,

383–426.

Spence, L.D., de Rosen-Spence, A.F., 1975. The place of sulphide

mineralization in the volcanic sequence at Noranda, Quebec. Economic

Geology 70, 90–101.

Staudigel, H., Muehlenbachs, K., Richardson, S.H., Hart, S.R., 1981.

Agents of low temperature ocean crust alteration. Contributions to

Mineralogy and Petrology 77, 150–157.

Staudigel, H., Davies, G.R., Hart, S.R., Marchant, K.M., Smith, B.M., 1995.

Large scale isotopic Sr, Nd and O isotopic anatomy of altered oceanic

crust: DSDP/ODP sites 417/418. Earth and Planetary Science Letters

130, 169–185.

Suzuoki, T., Epstein, S., 1976. Hydrogen isotope fractionation between

OH-bearing minerals and water. Geochimica et Cosmochimica Acta 40,

1229–1240.

Taylor, J.r., H, P., 1968. The oxygen isotope geochemistry of igneous rocks.

Contributions to Mineralogy and Petrology 19, 1–17.

Taylor Jr.., H.P., Sheppard, S.M.F., 1986. In: Valley, J.W., Taylor Jr.., H.P.,

O’Neil, J.R. (Eds.), Igneous rocks: I. Processes of isotopic fractionation

and isotope systematics Stable Isotopes in High Temperature

Geological Processes. Mineralogical Society of America, Reviews in

Mineralogy and Geochemistry, 16, pp. 227–271.

A. Perez-Aguilar et al. / Journal of South American Earth Sciences 18 (2005) 305–321 321

Vasquez, R., Vennemann, T.W., Kesler, S.E., Russell, N., 1998. Carbon and

oxygen isotope halos in the host limestone El Mochito Zn–Pb–(Ag) skarn

massive sulfide-oxide deposit Honduras. Economic Geology 93, 15–31.

Waring, C.L., Andrew, A.S., Ewers, G.R., 1998. Use of O C, and S stable

isotopes in regional mineral exploration. AGSO. Journal of Australian

Geology and Geophysics 17, 301–313.

Zengquian, H., Zaw, K., Xiaoming, Q., Qingtong, Y., Jinjie, Y., Mingji, X.,

Deming, F., Xianke, Y., 2001. Origin of the Gacun volcanic-hosted

massive sulfide deposit in Sichuan China: Fluid inclusion and oxygen

isotope evidence. Economic Geology 96, 1491–1512.

Zheng, Y-F., 1993a. Calculation of oxygen isotope fractionation in

anhydrous silicate minerals. Geochimica et Cosmochimica Acta 57,

1079–1091.

Zheng, Y-F., 1993b. Calculation of oxygen isotope fractionation in

hydroxyl-bearing silicates. Earth and Planetary Science Letters 120,

247–263.