485

Geological Characteristics of Epithermal Precious and Base Metal Deposits

STUART F. SIMMONS,†

Geology Department, University of Auckland, Private Bag 92019, Auckland, New Zealand

NOEL C. WHITE, P.O. Box 5181, Kenmore East, Queensland, Australia 4069

AND DAVID A. JOHN

U.S. Geological Survey, 345 Middlefield Rd., Menlo Park, California 94025

AbstractEpithermal deposits are important sources of gold and silver that form at <1.5-km depth and <300°C in

high-temperature, mainly subaerial hydrothermal systems. Such hydrothermal systems commonly develop inassociation with calc-alkaline to alkaline magmatism, in volcanic arcs at convergent plate margins, as well as inintra-arc, back-arc, and postcollisional rift settings. Many important deposits are Tertiary and younger in ageand are concentrated around the Pacific Rim and in the Mediterranean and Carpathian regions of Europe.Older deposits occur in the Tethyan arc from Europe to Asia and others are scattered in volcanic arcs of all ageswith rare examples as old as Archean.

Precious metal mineralization develops in zones of high paleopermeability, hosted within sequences of co-eval volcanic and underlying basement rocks. Veins with steep dips are common and these tend to host high-est grade ores. Precious metal mineralization also occurs in breccias, coarse clastic rocks, and intensely leachedrocks; such disseminated ore is much lower in grade but greater in total tonnage and may be amenable to bulkmining methods. Deposits and districts, comprising one or more orebodies, cover areas from <10 to ~200 km2.

Epithermal deposits have been classified on the basis of alteration and gangue mineral assemblages, metalcontents, sulfide contents, and sulfide mineral assemblages, and each classification scheme has its merits. Be-cause ores are oxidized by weathering, we prefer a classification that utilizes gangue mineral assemblages. Wedescribe two types of mineralization associated with quartz ± calcite ± adularia ± illite and quartz + alunite ±pyrophyllite ± dickite ± kaolinite assemblages, which reflect the pH of hydrothermal solutions.

Epithermal deposits associated with quartz ± calcite ± adularia ± illite contain Au-Ag, Ag-Au, or Ag-Pb-Znores. Electrum, acanthite, silver sulfosalts, silver selenides, and Au-Ag tellurides are the main gold- and silver-bearing minerals, with generally minor sphalerite, galena, and chalcopyrite; in some deposits base metals dom-inate the metal assemblage. Quartz is the principal gangue mineral accompanied by variable amounts of chal-cedony, adularia, illite, pyrite, calcite, and/or rhodochrosite, the latter in more Ag- and base metal-rich deposits.Distinctively banded crustiform-colloform textures, and lattice textures comprising aggregates of platy calciteand their quartz pseudomorphs, are common. Hydrothermal alteration is zoned and comprises deep regionalpropylitic alteration, which gives way upward to increasing amounts of clay, carbonate, and zeolite minerals,whereas quartz, adularia, illite, and pyrite form proximal alteration zones enveloping orebodies. Ore-grademineralization commonly terminates upward, and where there has been minimal erosion, it can be concealedbeneath regionally extensive blankets of clay-carbonate-pyrite or kaolinite-alunite-opal ± pyrite alteration.Fluid inclusion data indicate salinities are commonly <5 wt percent NaCl equiv for Au-Ag deposits and <10 to>20 wt percent NaCl equiv for Ag-Pb-Zn deposits. Stable isotope data indicate that hydrothermal solutionswere composed mostly of deeply circulated meteoric water, with a nil to small and variable component of mag-matic water.

Epithermal deposits associated with quartz + alunite ± pyrophyllite ± dickite ± kaolinite assemblages con-tain Au ± Ag ± Cu ores. Native gold and electrum are the main ore-bearing minerals, with variable amounts ofpyrite, Cu-bearing sulfides and sulfosalts such as enargite, luzonite, covellite, tetrahedrite, and tennantite, plussphalerite and telluride minerals; enargite dominates the Cu sulfides and indicates a high-sulfidation state.Quartz (both massive and vuggy) and alunite are the main gangue minerals with kandite minerals (dickiteand/or kaolinite) and/or pyrophyllite. Concentric patterns of hydrothermal alteration envelop the zone of vuggyand massive quartz alteration, which hosts ore. Outward, these comprise zones of quartz and alunite, dickite ±kaolinite or pyrophyllite, and illite or smectite alteration, surrounded by regional propylitic alteration. Zones ofillite or pyrophyllite alteration occur in the roots beneath some deposits. Fluid inclusion data indicate that salin-ities are typically <5 to 10 wt percent NaCl equiv but may be as high as >30 wt percent NaCl equiv. Stable iso-tope data indicate that the altering fluids are composed mostly of magmatic fluids with a minor to moderatecomponent of meteoric water.

Critical genetic factors include: (1) at several-kilometers depth, the development of oxidized and acidic ver-sus reduced and near-neutral pH solutions, controlled by the proportions of magmatic and meteoric compo-nents in solution, and the amount of subsequent water-rock interaction during ascent to the epithermal envi-

† Corresponding author: e-mail, (sf.simmons@auckland.ac.nz)

©2005 Society of Economic Geologists, Inc.Economic Geology 100th Anniversary Volumepp. 485–522

Introduction

EPITHERMAL deposits form in the shallow parts of high-tem-perature hydrothermal systems that commonly develop involcanic arcs (Fig. 1). The deposits are host to both preciousand base metals, but in the past three decades, they havebeen mined mainly for their gold and silver contents. Thetotal metal contents of some orebodies are substantial, and lo-cally the precious metal concentrations of some achieve bo-nanza grades (>1 Moz Au at >30 g/t; Sillitoe, 1993a). Some

deposits have been amenable to mining by simple methodsdating back many centuries (e.g., Abbot and Wolfe, 2003).The Spanish empire reached prominence during the colonialperiod (ca. 1500–1800 AD) through exploitation of the ep-ithermal ores of Mexico, Peru, and Bolivia, rich in either goldor silver. In the mid 1800s to early 1900s, epithermal discov-eries fueled gold-silver rushes to Nevada and New Zealand.During the past few decades, improved recovery methodsand favorable gold and silver prices (since the late 1970s)have enabled many low-grade orebodies to be mined. In total,

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ronment; (2) at epithermal depths, the development of boiling and/or mixing conditions which create sharpphysical and chemical gradients conducive to precious and base metal precipitation; and (3) at shallow level,the position of the water table, which controls the hydrostatic pressure-temperature gradients at depth whereepithermal mineralization forms.

Epithermal mineralization can occur in large areas, with orebodies that range in shape, size, and grade, andlie easily concealed beneath blankets of clay alteration or unaltered volcanic deposits. Efficient exploration re-quires integration of all geological, geochemical, and geophysical data, from regional to deposit scale. Vein min-eralogy and texture, patterns of hydrothermal alteration, patterns of geochemical dispersion, and three-di-mensional interpretation of related geophysical signatures are important guides. Willingness to drill is crucial,as surface features may not reliably indicate what is present at depth.

2 km

2 km

volcanic rocks

basement

intrusion

meteoric water

water table magmatic fluid

epithermal deposit

A Magmatic-Hydrothermal

100°200°

300°

BPD

250°

B Geothermal

vv

vv v

vvv

v

vv v

v

vv

v

vvv v

vvv

v

vv v

v

=

=

=

=

=

=

=

FIG. 1. Simplified conceptual models of high-temperature hydrothermal systems, showing the relationship between ep-ithermal environments, magmatic intrusions, fluid circulation paths, and volcanic and basement host rocks. A. The epither-mal environment forms in a magmatic-hydrothermal system dominated by acid hydrothermal fluids, where there is a strongflux of magmatic liquid and vapor, containing H2O, CO2, HCl, H2S, and SO2, with variable input from local meteoric water.This type of environment is analogous to those existing in modern volcanoes (e.g., Hedenquist et al., 1993; Christenson andWood, 1993). B. The epithermal environment forms in a geothermal system dominated by near-neutral pH chloride waters,where there is a strong flux of deeply circulated water (mostly of meteoric origin), containing CO2, NaCl, and H2S. This typeof system is analogous to those exploited for generation of electricity (e.g., Simmons and Browne, 2000a, b). The inferred lo-cation of the underlying magma chambers in both (A) and (B) are portrayed to show the different path lengths that deep flu-ids traverse before encountering the ore-forming environment. The relatively short path to the epithermal environment in(A) means there is minimal water-rock interaction during ascent, whereas the relatively long path to the epithermal envi-ronment in (B) means there is considerable water-rock interaction during ascent. The maximum pressure-temperature gra-dient under hydrostatic conditions is represented by boiling point for depth (BPD) temperatures, which are also shown forreference.

about 6 percent of all gold and about 16 percent of all silvermined have come from epithermal deposits (Singer, 1995),and their wide range of tonnage-grade characteristics(Hedenquist et al., 2000) make them an attractive target forboth large and small exploration and mining companies.

The term epithermal derives from the genetic classificationscheme for hydrothermal ore deposits proposed by Lindgren(1933). On the basis of stratigraphic relationships in volcanicsequences, and by analogy with mineral and metal occur-rences and mineral textures in active hydrothermal systems,Lindgren inferred that epithermal deposits formed at <200°Cand <100 atmospheres (~100 bars). Aside from the seminalpaper by D.E. White (1955) that strengthened the link to anactive hydrothermal environment, there was little advance inthe understanding of epithermal deposits until the late 1970swhen exploration interest rose due to the increasing value ofgold and silver. New research techniques applied to these de-posits included fluid inclusion studies that extended the rangeof formation temperature to about 300°C (e.g., Nash, 1972;Casadevall and Ohmoto, 1977; Kamilli and Ohmoto, 1977;Sawkins et al., 1979; Buchanan, 1981), and stable isotopestudies that indicated the prevalence of meteoric waters inthe formation of gangue minerals from some epithermal de-posits (e.g., O’Neil and Silberman, 1974; Casadevall andOhmoto, 1977; Kamilli and Ohmoto, 1977; Sawkins et al.,1979). Experimental and theoretical techniques were used todetermine metal solubilities and mineral stabilities under hy-drothermal conditions (e.g., Seward, 1973; Barton et al.,1977; Barnes, 1979), which led to numerical simulations ofreaction paths and ore formation (Reed, 1982; Drummondand Ohmoto, 1985; Reed and Spycher, 1985; Spycher andReed, 1989).

Meanwhile, in New Zealand, Japan, Philippines, UnitedStates, and other countries, the demand for alternativesources of electricity encouraged geothermal explorationdrilling and development. Temperatures and pressures simi-lar to those in the epithermal environment were encounteredat depths of less than 1 km (e.g., White, 1981; Henley andEllis, 1983), and precious and base metals were found de-posited in springs, wells, and surface pipes (e.g., Weissberg,1969, Hedenquist and Henley, 1985a; Brown, 1986; Kruppand Seward, 1987). The rapid increase in understanding atthe time was such that the first two volumes of Reviews inEconomic Geology focused on the nature of epithermal envi-ronments (Henley et al., 1984; Berger and Bethke, 1985).Thus, by the mid 1980s, genetic models were formulated toexplain the occurrence and zonation of metals and minerals,to define the physical-chemical conditions of ore depositionin several epithermal deposits, and to provide a basis for spec-ulation on the sources of fluids and metals (e.g., Barton et al.,1977; Kamilli and Ohmoto, 1977; Sawkins et al., 1979;Buchanan, 1981; Berger and Eimon, 1983; Henley and Ellis,1983; Hayba et al., 1985; Heald et al., 1987; Stoffregen,1987). In these models, hydrology was seen to be an essentialfactor in producing ore deposits, with boiling and fluid mixingbeing recognized as causative agents for metal deposition.Because they overlap in temperature and metal suite, Carlin-type deposits were initially included in the epithermal realmby several workers (e.g., Radtke et al., 1980; Berger andBethke, 1985; Radtke, 1985; Berger and Henley, 1989), but

they were later defined as a distinct class of sedimentary rock-hosted hydrothermal ore deposits (Kuehn and Rose, 1992,1995; Hofstra and Cline, 2000; Cline et al., 2005). The mod-ern use of the term epithermal thus retains much of Lind-gren’s intent and insight.

Since 1990, numerous articles have reviewed the natureand genesis of epithermal gold-silver deposits (e.g., Whiteand Hedenquist, 1990; Sillitoe, 1993a, b; Arribas, 1995;Richards, 1995; Simmons, 1995; Cooke and Simmons, 2000;Jensen and Barton, 2000; Sillitoe and Hedenquist, 2003). Inthis paper, we draw heavily on these references together withdata, mostly published since 1975, tabulated for more than 70epithermal deposits (see App. Table A1) that were selected torepresent the range of typical characteristics and geographicdistribution. The deposits cited below as examples are listedin the Appendix along with their relevant references, andtheir locations are shown in Figure 2. Plan maps of some de-posits (Fig. 3) show the main geologic features and the di-mensions of ore zones. Our aim in this paper is to provide anoverview of the diversity of features that characterizes ep-ithermal deposits and to relate these to common ore-formingprocesses and to suggest strategies for exploration.

Definition and ClassificationThe term epithermal refers to a range of temperature ver-

sus depth (pressure) ore-forming conditions that developwithin much larger, mainly subaerial, hydrothermal systems(Fig. 1). Depth relates directly to pressure in the shallow en-vironment where near-hydrostatic conditions prevail, withmaximum temperature largely controlled by the boiling-point-for-depth curve (e.g., Haas, 1971; Fig. 4). Ore mineralsprecipitate at temperatures ranging from ~150° to ~300°Cand at depths ranging from ~50 to as much as 1,500 m belowthe water table, caused by chemical changes that result fromsharp pressure and temperature gradients in this environ-ment. These physical controls define the epithermal environ-ment, although ore genesis also depends on the compositionof the hydrothermal solutions, which controls metal transportand deposition (e.g., Henley, 1985). Such metal-transportingsolutions vary in composition and differ in origin (e.g., Ar-ribas, 1995; Simmons, 1995) and thus vary in their metal en-dowment (e.g., Albinson et al., 2001).

Lindgren (1933) showed that despite sharing commongangue mineral assemblages, metal inventories of epithermaldeposits range widely, with varying proportions of gold, silver,and base metals, including mercury, antimony, tellurium, andselenium. Lindgren recognized their diverse characteristicswhen he distinguished nine deposit types based on metal con-tents (i.e., cinnabar, stibnite, base metal, gold, argentite gold,argentite, gold telluride, gold telluride with alunite, gold se-lenide). Nevertheless, epithermal deposits mined today areprincipally a source of precious metals.

Since the late 1970s, over a dozen classification schemeshave been proposed (Table 1). All of them consider some as-pect of ore or gangue mineralogy and most reflect some as-pect of the fluid chemistry (pH, oxidation state, or sulfidationstate) associated with proximal hydrothermal alterationand/or ore mineralization (Table 2). That so many schemeshave been proposed reflects the wide range of characteristicfeatures displayed by orebodies, as well as the evolution in

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thinking. Acid-base and reduction-oxidation fluid-mineralequilibria underpin the parameters that distinguished acidfrom alkaline types (Sillitoe, 1977), acid-sulfate or alunite-kaolinite from adularia-sericite types (Hayba et al., 1985;

Heald et al., 1987; Berger and Henley, 1989), and high-from low-sulfidation types (Hedenquist, 1987; White andHedenquist, 1990, 1995; Sillitoe, 1993a; White and Poizat,1995).

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Ma

La

Em

CV

EPEI-P

Ya

HB

CR

Po

Ke

Mc

TaPa

Ro

Fu

CrPj

Te

CCRM

Es

LC

Mi

Ch FrOv

CPJu

BM

Bo

Che

Hi

Le-ViGtoBa

PV

Pi

FIG. 2. Location of epithermal deposits listed in Appendix Table A1. Labels include most deposits mentioned in the text,but some are left out for clarity. Abbreviations: Ba = Baguio district (Acupan); BM = Baia Mare; Bo = Boliden; CC = Crip-ple Creek; Ch = Chinkuashih; Che = Chelopech; CP = Cerro de Pasco and Colquijirca-San Gregorio; Cr = Cracow; CR =Cerro Rico; CV = Cerro Vanguardia; EI-P = El Indio-Pascua; Em = Emperor; EP = El Peñon; Es = Esquel; Fr = Fresnillo;Fu = Furtei; Gto = Guanajuato; HB = Hope Brook; Hi = Hishikari; Ju = Julcani; Ke = Kelian; La = Ladolam; Le-Vi = Lep-anto-Victoria; LC = La Coipa; Ma = Martha Hill-Favona; Mc = McLaughlin; Mi = Misima; Ov = Ovacik; Pa = Pachuca-Realdel Monte; Pi = Pierina; Pj = Pajingo; Po = Porgera; PV = Pueblo Viejo; RM = Round Mountain; Ro = Rodalquilar; Ta =Tayoltita; Te = Temora; Ya = Yanacocha.

TABLE 1. Evolution of Classification Schemes Applied to Epithermal Deposits (modified from Sillitoe and Hedenquist, 2003)

Sillitoe (1977) Acid Alkaline

Buchanan (1981) Epithermal

Ashley (1982) Enargite gold

Giles and Nelson (1982) Hot-spring type

Bonham (1986, 1988) High sulfur Low sulfur Alkalic deposits

Hayba et al. (1985) Acid sulfate Adularia-sericiteHeald et al. (1987)

Hedenquist (1987), White and High sulfidation Low sulfidationHedenquist (1990, 1995)

Berger and Henley (1989) Alunite-kaolinite Adularia-sericite

Albino and Margolis (1991) Type 1 adularia-sericite Type 2 adularia-sericite

Sillitoe (1989, 1993a) High sulfidation Low sulfidation

High sulfide + base metal Low sulfide + base metal

White and Poizat (1995) High sulfidation Low sulfidation

Au-Ag-Cu Au-Ag-Cu Sn-Ag-base Ag-Au-base Au-Ag depositsdeposits with deposits with metal metal With calc-alkaline With alkaline vuggy quartz pyrophyllite-sericite deposits deposits volcanic rocks volcanic rocks

alteration alteration

Hedenquist et al. (2000), High sulfidation Intermediate sulfidation Low sulfidation AlkalicEinaudi et al. (2003),Sillitoe and Hedenquist (2003)

Cooke and Deyell (2003) Descriptive nomenclature based on ore metals, deposit form, diagnostic hypogene gangue and alteration minerals, and dominant Cu-bearing mineral

Application of the term “sulfidation,” which in respect toepithermal deposits was initially used to describe the oxida-tion state of aqueous sulfur species of deep ore-forming solu-tions (Hedenquist, 1987; Hedenquist and Lowenstern, 1994),was merged to agree with its other widespread use in orepetrology to describe the stabilities of sulfur-bearing mineralsin terms of sulfur fugacity (e.g., Barton and Skinner, 1967,1979; Hedenquist et al., 1994; Einaudi et al., 2003). This re-sulted from recognition that epithermal ore mineral assem-blages could be distinguished in terms of their high-, inter-mediate-, or low-sulfidation state (John et al., 1999) and thatfluids forming these assemblages could change sulfidationstates in response to chemical evolution both in space andtime (Einaudi et al., 2003). The variability in ore mineralogy(especially Fe-, Cu-, and As-bearing sulfides) and in the sulfi-dation states can be correlated to processes within the ep-ithermal environment, as well as to igneous rock compositionsand tectonic setting, the latter reflecting fundamental con-trols beneath the ore-forming environment (John et al., 1999;John, 2001; Sillitoe and Hedenquist, 2003; see Table 3). Al-though there is much remaining to learn about these rela-tionships, the sulfidation-state terminology reflects the evolu-tion of ascending hydrothermal fluids and assists inunderstanding the genesis of epithermal deposits (Einaudi etal., 2003).

The classification schemes in Table 1 that are based on al-teration and gangue minerals associated with gold and silver

ore are useful, especially at the early stages of prospect eval-uation, because their respective rock textures and alterationzonation patterns can be distinguished in the field or withminimal petrographic study. By comparison, the sulfide min-erals in epithermal deposits that occur near the surface and inthe vadose zone are susceptible to rapid oxidation and con-version to supergene minerals. Thus, the potential insights re-sulting from determining the sulfidation states of preciousmetal mineralization may be elusive or difficult to determinefrom field examination of rocks and may not be precisely es-tablished until exploration of a prospect or deposit is well un-derway. The two end-member types described below arebased on the hypogene gangue mineral assemblages that con-tain quartz ± calcite ± adularia ± illite and quartz + alunite ±pyrophyllite ± dickite ± kaolinite. These mineral assemblagesform from solutions of near-neutral and acid pH, respectively,but as discussed later in this paper, the fluid compositions in-ferred from these mineral assemblages may differ from thecompositions of ore-forming fluids transporting metals.Cooke and Deyell (2003) suggested another means of classi-fying deposits that is based on the metal contents, depositform, diagnostic hypogene gangue and alteration minerals,and the dominant Cu-bearing mineral (Table 1), similar to thedescriptive proposal made earlier by White and Hedenquist(1990). This classification has merit but is dependent on de-posit familiarity; the length of the names may limit its futureuse.

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TABLE 2. Diagnostic Minerals and Textures of Various States of pH, Sulfidation and Oxidation State Used to Distinguish Epithermal Ore-Forming Environments (Giggenbach, 1997; Einaudi et al., 2003)

(the use of hyphens between minerals indicates an equilibrium assemblage for which all phases need to be present)

Acid pH Neutral pHAlunite, kaolinite (dickite), pyrophyllite, Quartz-adularia ± illite, calciteresidual,vuggy quartz

High sulfidation Intermediate sulfidation Low sulfidationPyrite-enargite,± luzonite, covellite- Tennantite, tetrahedrite, hematite-pyrite- Arsenopyrite-loellingite-pyrrhotite, digenite, famatinite, orpiment magnetite, pyrite, chalcopyrite, pyrrhotite, Fe-rich sphalerite-pyrite

Fe-poor sphalerite-pyrite

Oxidized ReducedAlunite, hematite-magnetite Magnetite-pyrite-pyrrhotite, chlorite-pyrite

TABLE 3. Summary of Relationships between Sulfidation State of Ore-Forming Environment, Related Igneous Rock Compositions, and Tectonic Setting Proposed by Sillitoe and Hedenquist (2003)

Sulfidation state Igneous rock composition1 Tectonic setting

High Calc-alkaline, andesite-dacite Magmatic arc in a neutral to mildly extensional stress state; compressive stress state uncommon but serves to suppress volcanic activity

Intermediate Calc-alkaline, andesite-rhyolite Magmatic arc in a neutral to mildly extensional stress state; compressive stress state rare

Low Calc-alkaline, alkaline, tholeiitic bimodal basalt-rhyolite Magmatic arc undergoing extension leading to rifting; postcollisional rifting

1 Genetic relationship inferred by temporal-spatial correlation

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Vanguardia vein

O. D

iez vein

Atila vein

Luciana vein

Tres Patas vein

Vein

Cerro Vanguardia Guanajuato

Sierra vein system

La Luz vein system

Veta Madre vein system

131 t Au1,605 t Ag

Fault

Mine shaft

10 km010 km0

Cun

cuna

vei

n

Concep

cion vein

Lag. Mineral vein Fault

175 t Au34,850 t Ag

Rhyolite dikeQuartz vein

Fault

Klondyke

Rose’s Pride

Golden Plateau Golden Mile

Dawn

Pyropylitic

Unaltered rock

Calderamargin

Advanced argillic

Diorite

Vuggy silica

Ore zone

Monte Negro Moore

Intermediate argillic

Rodalquilar

Abundant Stage 2 alunite

Pueblo Viejo

Cracow Rodalquilar

4 km0 4 km0

4 km0

Central Extended

1,242 t Au7,062 t Ag

10 t Au26 t Au30 t Ag

Guinaoang

Lepanto

Victoria epithermaldeposit

115 t Au393 t Ag

VeinCu-Au mineralization

Basement rocksPorphyry Cu-Au

Fault

4 km0

Lepanto epithermal Cu-Au deposit

FSE porphyry Cu-Au

TonaliteDacite

FIG. 3. Sketch maps of epithermal deposits, showing the outlines of orebodies grouped according to scale. These illus-trate the great variability in the sizes and shapes of orebodies. Note that total production correlates poorly with the areal ex-tent of ores, which is further reflected in the data set reported in Appendix Table A1. Maps are redrawn as presented in pub-lications so there is some inconsistency in, for example, locating the occurrences of veins. References for each deposit aregiven in the Table A1.

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Quartz Monzonite Dike

Vein

Esperanza vein

2270 vein

Caldera contact

Caldera contact (360m below surface)

DikeMajor Fault

Flatmake (vein)

Fresnillo

Santa Elena vein

Santo Nino vein

San Ricardo veinSan Carlos vein

San Miguel vein Santa Cruz vein

San Emetrio veinSan Mateo vein

Santa Paula veinSanta Inocencia vein

2 km0 2 km0

EmperorFresnillo 136 t Au26 t Au16,050 t Ag

Empire

Royal

Martha

Edw

ard

Martha HillOpen pitoutline

Cowshed

Union

Silverton Fa

vona

Old Favona

AmaranthMascotte

Gladst

one

Moo

nlig

ht

Sanjin Ore Zone

Honko Ore Zone

Yamada Ore Zone

Vein

Balatoc diatreme

Fault

1 km0

1 km0 1 km0

Acupan Martha Hill-Favona

Hishikari 260 t Au140 t Ag

263 t Au1,253 t Ag

200 t Au200 t Ag

Jalene

El Indio

Campana

La Vieja

Viento

Bechita-HuantinaCanto Norte

VeinFault

1 km0 1 km0

Underground stopes and ore zones with Au >10 g/t

Au >0.34 g/t

Fault

Open pit outline

Missionary vein

El Indio Summitville 17 t Au23 t Ag

310 t Au3,100 t Ag

FIG. 3. (Cont.)

General Characteristics of Epithermal DepositsEpithermal deposits comprise epigenetic ores that are gen-

erally hosted by coeval and older volcanic rocks and/or un-derlying basement rocks and rarely by subvolcanic intrusions.They cover areas that range from <10 to >100 km2 (Fig. 3).The orebodies occur in a diversity of shapes that reflect theinfluence of structural and lithological controls, and they rep-resent zones of paleopermeability within the shallow parts ofonce active hydrothermal systems (Figs. 3, 5). Most com-monly, orebodies occur in veins with steep dips that formedthrough dilation and extension. Some are hosted by majorfaults but more commonly they are hosted by minor faults(second- or third-order structures) with small displacements(<10 m). Optimum structural development generally de-pends on rock rheology and brittle failure. Lithology is alsoimportant, especially where contrasts in porosity and perme-ability focus the fluid flow through specific units, along rockcontacts, or through permeable masses of brecciated rock.These lithologic features may be an intrinsic characteristic ofthe original rock; alternatively, they may be a by-product ofhydrothermal alteration and chemical dissolution or hy-drothermal brecciation (Sillitoe, 1993b). Thus, faults andfracture networks, as well as breccias, coarse clastic rocks, and

intensely leached rocks account for the spectrum of vein-re-lated to disseminated ores (Table 4), which can extend for100s to 1,000s of meters laterally and 10s to 100s of metersvertically. The dominant gangue mineral is quartz, makingores hard and generally resistant to weathering, and the dom-inant sulfide mineral is pyrite, with sulfide contents that canrange from <1 to >20 vol percent.

Metal endowments

Most ores are mined for gold and silver, and there is a spec-trum of gold-rich (Ag/Au ratio <10, locally <1; e.g., Cracowand Pajingo, Australia; Hishikari, Japan; Acupan and Antamokin Baguio, Philippines; Ladolam, Papua New Guinea; RoundMountain, United States) to silver-rich deposits (Ag/Au ratios~20–200; e.g., Pascua-Lama and La Coipa, Chile; Tayoltita,Guanajuato, and Pachuca-Real del Monte, Mexico; Comstockand Tonopah, United States). Some of these are copper bear-ing, with high- to intermediate-sulfidation–state mineral as-semblages containing As and Sb (e.g., Yanacocha, Peru; ElIndio, Chile; Lepanto, Philippines; Goldfield, United States).There are also Ag-Pb-Zn deposits (with subordinate Cu, As,and Sb) that are poor in Au (Ag/Au ratio >400), and their dis-tribution is restricted to provincial belts of mineralization, best

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kaol

inite

kaol

inite

Boiling p

oint for dep

th curve (pure w

ater)

epid

ote

epid

ote

bio

tite

wai

raki

tela

umon

tite

illite

& c

hlor

ite

illite

inte

rlaye

red

cla

ys

inte

rlaye

red

cla

ys massive opal

massive opal

cristobalite + sulfur

alun

ite

smec

tite

kaol

inite

acid alteration

sinter

and

alus

ite

300

300

200

200

100

100

0

1000

500

A B C

Dd

iasp

ore

pyr

ophy

llite

dic

kite

smec

tite

smec

titeTemp (°C)

Temperature (°C)

Depth (m)

Alteration zonation atthe water table

watertable

watertable

40

20

0

Depth (m)

bio

tite

Vertical distributionof minerals

in boiling upflow zone

FIG. 4. Key indicator minerals in epithermal environments. A. Stability range of temperature-sensitive clays, phyllosili-cates, and zeolites (Henley and Ellis, 1983; Reyes, 1990). B. Vertical distribution of some of the same minerals plotted ac-cording to depth, using the hydrostatic boiling curve as the reference temperature gradient. C. Diagnostic hydrothermalminerals forming at the water table, comprising silica sinter where near-neutral pH waters discharge around boiling hotsprings and vertically zoned acid alteration (modified from Sillitoe, 1993b). D. Magnification of vertically zoned steam-heated acid alteration at the water table (Schoen et al., 1974; Simmons and Browne, 2000a): cristobalite and sulfur form atand above the water table; tabular massive opal forms at and below the water table; alunite and kaolinite form at and belowthe water table and the zone of massive opal.

exemplified by those occurring in northern Mexico (e.g., Fres-nillo and Zacatecas). Even more isolated in occurrence are ex-amples like Cerro Rico de Potosí, Bolivia, which, albeit ep-ithermal in style and the largest silver deposit in the world, isa variant of mineralization found in the Ag-Sn belt of Bolivia,where deposits formed at conditions generally deeper and hot-ter than epithermal (e.g., Sillitoe et al., 1998).

Geologic parameters affecting production of epithermal oresinclude mineralization style (e.g., structurally controlled versusdisseminated), grade distribution, and supergene oxidation, allof which can affect the cost of mining and metallurgical pro-cessing. Open-pit methods are used for large tonnage low-grade (1–2 g Au/t; 90 g Ag/t) orebodies (e.g., Round Mountain,United States; Real de Angeles, Mexico), with gold in oxidized

ores being amenable to low-cost heap-leach treatment. Under-ground methods are generally used to exploit small to modesttonnage but high-grade (10–>100 g Au/t, >500 g Ag/t) orebod-ies (e.g., Hishikari, Japan; Emperor, Fiji; Fresnillo, Mexico), ex-cept where they intersect the surface and can be mined fromopen cuts (e.g., Cerro Vanguardia, Argentina). At finer scale,within individual orebodies, the highly variable nature of goldand silver assays over distances of less than a few meters, makesgrade control a critical part of successful mining, especiallywhere ores are hosted in structural zones.

Relationship to igneous rocks

Most epithermal deposits are affiliated with coeval volcanicrocks and subvolcanic intrusive equivalents of predominantly

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200 m0

Ignimbrite

Altered andesite

100

sealevel

-100

-300

-200

-400

NMartha Hill Ladolam

S

Martha vein

Roy

al v

ein

Drea

dnou

ght v

ein

Empi

re v

ein

Paleosol

NS

NS

1 - 3

Vent breccia

Gold Grade (g/t)

3 - 7

Porphyry breccia

Coherent volcanic rocks

7 - 15

Volcanic breccia

>15

Alkaline intrusions cut by porphyry breccia

500 m0

-300

-200

-100

sealevel

100

-300

-200

-100

sealevel

100

FIG. 5. Examples of structural and lithological controls on orebody geometry. At Martha Hill, precious metal mineraliza-tion is entirely hosted in steeply dipping veins extending over a 500-m vertical interval (redrawn from Morgan, 1924); notethat the tops of the veins are cut by an erosional unconformity overlain by unaltered ignimbrite. At Ladolam, disseminatedgold mostly occurs in subhorizontal tabular zones hosted by a complex sequence of volcanic and hydrothermal breccias (re-drawn from Carman, 2003).

TABLE 4. Geometrical Controls on Permeability and Epithermal Orebodies (modified from Sillitoe, 1993b)

Control Orebodies Permeability control Examples

Structural: Veins (steeply dipping); vein Extension-transtension second- and third-order Martha Hillinfluenced by faults swarms and stockworks; structures; dilational jogs and releasing bends; McLaughlinand fractures fault intersections brittle fracturing Tayoltita

HishikariThamesComstock

Hydrothermal: influenced Hydrothermal breccia; Overpressuring followed by hydraulic fracturing Ladolamby the pressure and diatremes; and residual or hydrothermal eruption; intense acid leaching Cripple Creekreactivity of fluids vuggy quartz Summitville

Nansatsu districtAcupanGoldfieldYanacochaPascua-Lama

Lithological: influenced Strata-bound disseminations Coarse-grained ignimbrite and/or sedimentary Round Mountainby the physical unit that is unconsolidated or that has easily Chinkuashihcharacteristics of rocks dissolved cement; rock contacts separating Yanacocha

permeable and impermeable strata

calc-alkaline affinity that form in magmatic arcs resultingfrom convergent plate movement and plate subduction(Sawkins, 1990; Sillitoe and Hedenquist, 2003). Gold-silver,Au ± Ag ± Cu, and Ag-Pb-Zn deposits are all found in vol-canic sequences containing andesite, dacite, and rhyolite.These calc-alkaline magmas are relatively oxidized (magmaticoxygen fugacity ≥ nickel-nickel oxide buffer; e.g., Hildreth,1981; John, 2001; Einaudi et al., 2003) and generated by par-tial melting of the mantle wedge above subducting oceaniclithosphere (e.g., Gill, 1981; Luhr, 1992). Epithermal Au-Agdeposits of relatively low Ag/Au ratio are also found with vol-canic rocks that erupted in back-arc and continental-rift envi-ronments, producing reduced tholeiitic magmas with bimodalbasalt-rhyolite compositions. The best documented examplesare in the Great Basin of the western United States (Hildreth,1981; John, 2001).

There are some important exceptions to these generaltrends, including the few, but very large, Au-Ag ± Te depositsthat are closely related to alkaline volcanic rocks that were de-rived from oxidized and hydrous mafic magmas (Richards,1995; Jensen and Barton, 2000). Such magmas form outsideconventional volcanic arcs in zones of crust where deeplypenetrating tensional structures developed through rifting(e.g., Cripple Creek, United States; Ladolam, Papua NewGuinea; Emperor, Fiji) or postsubduction tectonism (e.g.,Porgera, Papua New Guinea; Sillitoe, 1993a; Richards, 1995;Jensen and Barton, 2000). The correlation of magma compo-sition and metal assemblage is also seen at Cerro Rico de Po-tosí, where host volcanic rocks for the Ag-Sn ores consist ofrelatively reduced ilmenite-bearing rhyodacite (Sillitoe et al.,1998). The late Pliocene-Pleistocene age McLaughlin de-posit, California, formed during activity of the Clear Lake vol-canic field that erupted in response to upwelling of mantlethrough a slab window in a largely transpressional environ-ment east of the San Andreas transform fault (Sherlock et al.,1995; Dickinson, 1997). These exceptions highlight the widerange of tectonic settings that can host mineralization notedby Sillitoe and Hedenquist (2003).

Preservation in the geologic record

Given the relatively shallow depth of formation, epither-mal deposits may have poor preservation potential in the ge-ologic record, because they commonly form in high-reliefvolcanic arc settings and because convergent plate bound-aries are especially prone to phases of rapid uplift and ero-sion. Thus, a majority of deposits are Tertiary or younger(Table A1), and there are major deposits that have formedsince 2 Ma (e.g., Lepanto, Philippines; Hishikari, Japan;Ladolam, Papua New Guinea; McLaughlin, United States).However, older deposits have been preserved where theirhost volcanic belts are well preserved, such as the Mesozoicdeposits (e.g., Cerro Vanguardia, Argentina) of the Deseadomassif in Patagonia and the Paleozoic deposits (e.g., Temora,Pajingo, and Cracow) of the Tasman fold belt in eastern Aus-tralia, as well as similar examples in Mongolia and Russia(Yakubchuk et al., 2005). Precambrian examples are also re-ported for Canada, Scandinavia, and Australia but, to date,the known very ancient epithermal deposits are small (Dubéet al., 1998; Hallberg, 1994; Turner et al., 2001; Huston etal., 2002).

Active Epithermal EnvironmentsActive epithermal environments in geothermal and mag-

matic hydrothermal systems (Fig. 1) were important to theconception and classification of epithermal deposits (Ran-some, 1907; Lindgren, 1933). Such high-temperature hy-drothermal systems are located in geologic settings analogousto epithermal deposits (Henley and Ellis, 1983; Henley,1985), and they provide a context in which the mineral prod-ucts of hydrothermal activity can be compared with coexistingfluids at known temperatures, pressures, mass flows, andchemical compositions. For example, the occurrence of spec-tacular sulfide scales, containing 6 wt percent Au and 30 wtpercent Ag, on back-pressure plates (downstream of thethrottle point) within surface pipe work at the Broadlands-Ohaaki geothermal field was shown to be the direct conse-quence of boiling (flashing) of a fluid at 260° to 180°C initiallycontaining about 1 to 2 ppb Au (Brown, 1986). Although thelow-salinity (<0.5 wt % NaCl) and near-neutral pH solution isinitially undersaturated in gold and silver, the flashing envi-ronment results in quantitative precipitation of precious met-als, highlighting the efficiency of metal precipitation inducedby boiling in the epithermal regime. With geothermal wellsdrilled to >2.5-km depth (>300°C), such active systems pro-vide an overview of hydrothermal processes occurring within,above, below, and on the periphery of the epithermal envi-ronment (e.g., Henley and Ellis, 1983; Hedenquist, 1990;Reyes, 1990; Simmons and Browne, 2000a, b). Here webriefly examine the main fluid types and corresponding hy-drothermal mineral assemblages of active environments(Henley and Ellis, 1983; Giggenbach, 1992a, 1997) as aframework for understanding hydrothermal minerals in ep-ithermal deposits (Table 5), described in greater detail below.

Geothermal systems

Geothermal systems in volcanic arcs and rifts involve deepconvective circulation of meteoric water driven by shallow in-trusion of magma at >4-km (?) depth. At the deepest level ex-plored by geothermal wells, these chloride waters—so-calleddue to the dominant anion—are reduced and have near-neu-tral pH and contain from 0.1 to >1 wt percent Cl, up to 3 wtpercent CO2, and 10s to 100s of ppm H2S; the latter is an im-portant ligand for aqueous transport of gold and silver asbisulfide complexes (Seward, 1973; Seward and Barnes,1997). The concentrations of the main aqueous constituentsrepresent equilibrium with quartz, albite, adularia, illite, chlo-rite, pyrite, calcite, and epidote, which form as secondaryminerals during alteration of igneous rocks (Barton et al.,1977; Giggenbach, 1997). The fluid reaches equilibrium withthe rock and its constituent minerals where flow is slow,through a “rock-dominated” or rock-buffered environment,to form a propylitic alteration assemblage (Giggenbach,1997). Boiling occurs in the central upflowing column of fluiddown to 1- to 2-km depth below the water table, controlled bynear-hydrostatic pressure-temperature conditions (Fig. 4). Inthis environment, quartz, adularia, and calcite (usually platy)deposit in open spaces and subvertical channels from theboiling and cooling liquid (e.g., Simmons and Browne,2000b). Depending on the permeability structure, the chlo-ride water may rise to the surface to discharge and deposit

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silica sinter where topography intersects the geothermalwater table; alternatively, this liquid may disperse laterallythrough an outflow zone.

By contrast, dissolved gases (mainly CO2 and H2S) separatefrom the liquid into vapor due to boiling and rise to the sur-face along a path distinct from the residual liquid. The risinggases, CO2 and H2S, may be partially absorbed into coolground waters at shallow levels, along with condensed watervapor, to form two types of steam-heated waters, CO2-richand acid-sulfate. CO2-rich steam-heated waters contain highconcentrations of dissolved CO2 (>1 wt %) and tend to accu-mulate at shallow levels. They drape the stagnant margins ofthe upflow zone to depths as much as 1,000 m below thewater table. Their distribution is best known at Broadlands-Ohaaki, where weakly acidic steam-heated waters alter vol-canic rocks to an argillic assemblage dominated by clay min-erals (illite, illite-smectite, smectite, and kaolinite), calcite,and siderite at temperatures up to about 150°C (Hedenquist,1990; Simmons and Browne, 2000b).

Acid-sulfate steam-heated waters are close to 100°C andform in the vadose zone where H2S comes into atmosphericcontact and oxidizes to H2SO4. Their pH is ~2, and they con-tain relatively high concentrations of sulfate (~1,000 mg/kg).These waters alter rocks to an advanced argillic assemblageof opal (cristobalite), alunite, kaolinite, and pyrite as the so-lution is neutralized near the water table (Schoen et al.,1974). The distribution of these three water types largely de-pends upon topographically controlled hydraulic gradients.In low-relief volcanic settings (e.g., calderas, flow-domecomplexes, rifts), the steam-heated waters occur above andon the periphery of the chloride-water plume, whereas inhigh-relief settings (e.g., andesitic composite cones), thesteam-heated waters may extend from the summit to thelower flanks of the volcanic edifice; under the influence ofsuch a steep hydraulic gradient, chloride waters may flow lat-erally long distances (>5 km) to form subsurface outflowzones (Henley and Ellis, 1983). Hybrid compositions formwhere the waters mix.

Magmatic hydrothermal systems

Magmatic hydrothermal systems, unlike geothermal sys-tems, are rarely drilled because of their acidic conditionsand high temperatures. What we know of subsurface condi-tions is from gases discharged from fumaroles at 100° to>800°C, acidic hot springs, and hydrothermally alteredrocks ejected by explosive eruptions (e.g., Hedenquist et al.,1993). An exception is in the Philippines, where severalmagmatic hydrothermal systems with zones of very reactivefluids have been explored for their geothermal energy po-tential (Reyes, 1990; Delfin et al., 1992; Reyes et al., 1993,2003). Existing data on the metal contents of high-tempera-ture volcanic discharges indicate the potential for substan-tial flux of both precious and base metals (Hedenquist,1995). Within the central upflow column overlying shallowintrusions, the fluids in these systems are dominated bymagmatic components, including HCl, SO2, and HF. Whenthese gases condense into the hydrothermal system, SO2

disproportionates, forming H2S and H2SO4 (Sakai and Mat-subaya, 1977; Rye et al., 1992) and a very acidic (pH ~1) so-lution, containing subequal amounts of HCl and H2SO4, upto ~1 wt percent each (Giggenbach, 1997). Hydrolysis reac-tions with igneous country rocks progressively neutralizesthe acidity while forming hydrothermal minerals that in-clude alunite, pyrophyllite, dickite, quartz, anhydrite, dias-pore, and topaz, as well as kaolinite and illite, characteristicof “fluid-dominated” alteration conditions (Reyes, 1990;Giggenbach, 1992a, 1997). Surficial steam-heated acid-sul-fate waters also form in magmatic hydrothermal systems,just as they do in the vadose zone over geothermal systems,due to the presence of H2S in the vapor. Silica sinters, how-ever, are absent due to the acidic conditions that inhibit sil-ica polymerization and deposition of vitreous amorphous sil-ica (Fournier, 1985). In this setting, two styles of advancedargillic alteration, magmatic hydrothermal and steam-heated, develop with different origins both containing alu-nite and kaolinite (Rye et al., 1992).

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TABLE 5. Summary of Hydrothermal Alteration Assemblages Forming in Epithermal Environments

Alteration Mineralogy Occurrence and origin

Propylitic Quartz, K-feldspar (adularia), albite, illite, Develops at >240°C deep in the epithermal environment through chlorite, calcite, epidote, pyrite alteration by near-neutral pH waters

Argillic Illite, smectite, chlorite, inter-layered clays, Develops at <180°C on the periphery and in the shallow epithermal pyrite, calcite (siderite), chalcedony environment through alteration by steam-heated CO2-rich waters

Adv. Argillic (steam-heated) Opal, alunite (white, powdery, fine-grained, Develops at <120°C near the water table and in the shallowest pseudocubic), kaolinite, pyrite, marcasite epithermal environment through alteration by steam-heated

acid-sulfate waters; locally associated with silica sinter but only in geothermal systems

Adv. Argillic (magmatic Quartz, alunite (tabular), dickite, pyrophyllite, Develops at >200°C within the epithermal environment through hydrothermal) (diaspore, zunyite) alteration by magmatic-derived acidic waters

Adv. Argillic (supergene) Alunite, kaolinite, halloysite, jarosite, Fe oxides Develops at <40°C through weathering and oxidation of sulfide-bearing rocks

Advanced argillic alteration

The origin of advanced argillic alteration can be deter-mined from its morphology, as well as mineralogy and zona-tion (Table 5), and this information can be used to interpretthe level of exposure and proximity to potential epithermalmineralization (Sillitoe, 1993a; Hedenquist et al., 2000). Mag-matic hydrothermal or hypogene, advanced argillic alterationincludes minerals that form at >200°C, such as pyrophyllite,dickite, diaspore, zunyite, and topaz, with alunite that is gen-erally tabular and sometimes coarse grained. This alteration isepigenetic in nature, so it generally cuts across stratigraphyand follows high-angle structures, although it can be strati-form in permeable host rocks.

Steam-heated advanced argillic alteration forms above thewater table at ~100°C in horizons with pronounced verticalmineral zonation. In general, this blanket of alteration doesnot exceed 10 to 20 m in thickness. Tabular but discontinuousbodies of massive opal mimic and mark the water table, un-derlain by a discontinuous zone comprising alunite, kaolinite,opal, and variable amounts of pyrite and marcasite that givesway with depth to a kaolinite zone comprising kaolinite plusopal (Schoen et al., 1974; Simmons and Browne, 2000a; Fig.4). These alteration minerals are typically very fine grained,and the alunite generally occurs as pseudocubic crystals.

A third type of advanced argillic alteration is formed by su-pergene weathering and oxidation of sulfide-rich rocks thatpostdate hydrothermal activity. This alteration forms at<40°C, within the vadose zone, and comprises alunite, kaoli-nite, halloysite, jarosite, and iron oxides and hydroxides. Su-pergene advanced argillic alteration also has a blanketlikegeometry that mimics topography, but it may line sub-verticalfractures that were pathways for descending surface water.

A combination of careful geologic mapping and mineralidentification (with a hand lens, infrared spectrometer, petro-graphic microscope, X-ray diffraction, or scanning electronmicroscope) are generally sufficient for distinguishing the ori-gins of advanced argillic alteration. Rye et al. (1992) and Rye(2005) further describe how the alunite and kaolinite formingin these three environments can be distinguished on the basisof sulfur, oxygen, and hydrogen isotope analysis.

Mineralization Associated with Quartz ± Calcite ±Adularia ± Illite Assemblages

One type of epithermal mineralization is distinguished byits intimate association with quartz ± calcite ± adularia ± illitethat forms from the near-neutral pH chloride waters in ex-tinct geothermal systems. This gangue mineral assemblagehosts a spectrum of Au- to Ag-rich ores, as well as the Au-Ag± Te ores associated with alkaline rocks, and the Ag-Pb-Znores of northern Mexico. Quartz and/or chalcedony domi-nate, accompanied by lesser and variable amounts of adularia,calcite, pyrite, illite, chlorite and rhodochrosite; with the ex-ception of quartz, there are many examples where one ormore of these phases is missing or is trace in amount. Oresoccur in veins and stockworks, making up subvertical frac-tures, or, more rarely, in pore space of breccias and perme-able rocks, forming disseminated mineralization. In Au-Agdeposits, gold typically occurs as microscopic to submicro-scopic grains of electrum and rare tellurides, whereas silver

occurs as electrum, acanthite, sulfosalts (e.g., pyrargyrite-proustite, Ag-rich tetrahedrite) and/or silver selenide miner-als. Both precious metals are found with highly variableamounts of base metal sulfides (sphalerite, galena, and lesserchalcopyrite) and pyrite, marcasite, and/or pyrrhotite. Sul-fides constitute from <1 to >10 vol percent of the ore, and thesulfide abundance, particularly the base metal sulfides, insome deposits increase with depth or with changes in host-rock composition. Sulfidation states inferred from ore-relatedsulfide minerals range from intermediate to low (Heald et al.,1987; John, 2001; Einaudi et al., 2003; Sillitoe and Heden-quist, 2003).

Epithermal deposits are also distinguished by the ganguemineral textures (Fig. 6). Crustiform banded quartz is com-mon, typically with interbanded, discontinuous layers of sul-fide minerals (mainly pyrite) and/or selenide minerals, adu-laria, and/or illite. At relatively shallow depths, the bands arecolloform in texture and millimeter-scale, whereas at greaterdepths, the quartz becomes more coarsely crystalline. Latticetextures, comprised of platy calcite and its quartz pseudo-morphs, occur as open-space filling in veins, and along withvein adularia indicate boiling fluids of near-neutral to alkalinepH (Simmons and Christenson, 1994; Simmons and Browne2000b).

Breccias in veins and subvertical pipes commonly show ev-idence of multiple episodes of formation. They comprise jum-bled angular clasts of altered host rock and earlier vein fill,supported by a matrix of mainly quartz, calcite, and/or adu-laria and sulfide minerals (Fig. 6), suggesting rapid pressurerelease and violent formation that can be ascribed to seismic-ity (e.g., Sibson, 1987) and hydrothermal eruptions (e.g.,Hedenquist and Henley, 1985a).

Broad-scale patterns of alteration zoning surround orebod-ies and reflect the level of exposure (Fig. 7). At regional scale,deep level (>400 m below the water table) alteration is propy-litic (e.g., Acupan, Philippines; Comstock Lode and RoundMountain, United States; Tayoltita, Mexico; Martha Hill,New Zealand; cf. Hudson, 2003). At intermediate levels(400–150 m below the water table), clay and carbonate min-erals increase at the expense of aluminosilicate minerals,whereas zonation of clays (illite to smectite), and zeolites(wairakite to heulandite to mordenite) reflect decreasingtemperature (Fig. 4). Intense quartz, adularia, illite, andpyrite alteration commonly surrounds ores and reflects thesharp increase in permeability associated with fluid conduits;accordingly, in host rocks with low permeability, alterationmay be closely restricted to the selvages of veins and veinlets.At shallow levels (150–0 m below the water table), blankets ofargillic alteration, illite and other clays (with or without dis-seminated pyrite, carbonate, minor barite, and minor anhy-drite) are generally well developed, especially in host volcanicrocks, and may conceal underlying orebodies (e.g., Creede,United States; Pachuca-Real del Monte, Mexico; Barton etal., 1977; Dreier, 1982). At the shallowest depths in the ep-ithermal environment, steam-heated advanced argillic alter-ation occurs with or without silica sinters that form near thepaleowater table and the paleosurface (Figs. 4, 6). Silica sin-ter, which deposits as amorphous silica and then converts toquartz (Herdianita et al., 2000), shows rhythmic banding,plant fragments, and diagnostic columnar structures and may

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FIG. 6. Photographs of minerals and textures that commonly occur in epithermal deposits associated with quartz ± cal-cite ± adularia ± illite: A. Cinnabar-bearing silica sinter (Puhipuhi, New Zealand; scale bar = 2 cm). B. Colloform crustiformbanding in gold-silver–bearing ore (Martha Hill, New Zealand; scale bar = 2 cm). C. Adularia encrusted on open fracture(Martha Hill, New Zealand; scale bar = 1 cm). D. Lattice textures in which platy calcite is replaced by quartz in gold-sil-ver–bearing ore (Martha Hill, New Zealand; scale bar = 3 cm). E. Vein containing coarsely crystalline quartz, sphalerite, andgalena (Pachuca-Real del Monte, Mexico; scale bar = 1.25 cm). F. Brecciated vein material in gold-silver–bearing ore(Golden Cross, New Zealand; scale bar = 4 cm).

A B

C D

E F

be preserved in rock sequences containing epithermal de-posits (White et al., 1989).

Fluid inclusion data

Fluid inclusion studies, mostly on transparent ganguephases (quartz, calcite) and sphalerite (the main ore-relatedsulfide mineral suitable for fluid inclusion study), indicate oredeposition from dilute to moderately saline solutions at tem-peratures between 150° and 300°C. Gold-silver deposits gen-erally have dilute solutions of <5 wt percent NaCl equiv,whereas Ag-Pb-Zn deposits commonly have brines of <10 to>20 wt percent NaCl equiv (Fig. 8). Coexisting liquid- andvapor-rich fluid inclusions are common and indicate boilingconditions at the time of trapping (Bodnar et al., 1985). Thisallows temperatures of boiling to be used to calculate pres-sures and estimate depths of formation (Roedder and Bodnar,1980). Therefore, assuming a hydrostatic boiling-point-for-depth gradient (Haas, 1971), consistent with estimates of

vertical temperature gradients (Vikre, 1985; Simmons et al.,1988; Cooke and Bloom, 1990; Sherlock et al., 1995) and ge-othermal system analogues, ore deposition occurs over adepth range of about 50 to 1,100 m below the water table.These are minimum values, however, because the presence ofsmall amounts of dissolved CO2, the main gas in geothermalfluids (Hedenquist and Henley, 1985b), increases the totalfluid pressure by as much as several tens of bars and increasesthe depth range of boiling up to hundreds of meters (e.g.,Simmons, 1991; Sherlock et al., 1995).

Stable isotope data

Stable isotope studies, comprising measurements of δDand δ18O, have been made on several gangue minerals(quartz, adularia, clays, and carbonates) and on fluid inclu-sions to determine the provenance of the fluid responsible foralteration and mineralization; few of the studies have de-termined the isotopic composition of the ore solutions

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Quartz + Alunite ± Pyrophyllite ± Dickite ± Kaolinite

Quartz ± Calcite ± Adularia ± Illite

quartz, chalcedony, adularia, carbonatespyrite, Au-Ag, Ag-Pb-Zn

lattice textures, crustiform-colloform banding

quartz, illite, adularia, pyrite clay

carbonatepyrite

propylitic50

-100

m

50-100 m

vuggy to massive quartznative Au, sulfosalts, pyrite

quartz, alunite

1-10 m

propylitic

50-1

00 m

50-100 m

dickite (kaolinite)pyrophyllite

smectitemixed layer clay

1-10 m

propylitic

quartz, alunitedickite (kaolinite)

pyrophyllite, pyrite

FIG. 7. Sketch diagrams showing the mineralogic zonation at two different scales around epithermal orebodies associatedwith quartz ± calcite ± adularia ± illite and quartz + alunite ± pyrophyllite ± dickite ± kaolinite gangue mineral assemblages.The diagrams on the left show the large-scale pattern, and the rectangle area outlined is magnified on the right to show al-teration zonation patterns in the vicinity of ore (after Sillitoe, 1993b).

themselves (Fig. 9). The interpretation of such data is notstraightforward, because the data typically are scattered,water compositions generally have to be constructed fromanalyses of different minerals (hydroxyl-bearing clays) or fluidinclusion waters, and equilibration (or fractionation) temper-atures have to be estimated. In addition, doubt has been caston the validity of δD analyses of quartz-hosted fluid inclusionwaters, as they may yield unreliable values that are too low ifthe quartz crystallized from originally precipitated amor-phous silica or if the waters are extracted by thermal decrepi-tation (Faure et al., 2002; Faure, 2003). Deposits youngerthan a few million years generally allow more accurate con-straints on the composition of local meteoric water, with pre-sent-day values serving as a reliable proxy. Notwithstanding

these problems, the results generally plot between the mete-oric water line and compositions associated with magmaticwater (Fig. 9), suggesting that mixing of waters from bothsources accounts for the compositions measured (e.g., O’Neiland Silberman, 1974; Faure et al., 2002). Commonly, inter-pretations are inconclusive, because water-rock interaction ofdeeply circulated meteoric water results in an evolution ofisotopic compositions—the “18O-shift” (Craig, 1963; Taylor,1979). This overlap in isotopic compositions has caused con-siderable debate on the origins of waters in subaerial geot-hermal systems (e.g., Giggenbach, 1992b, 1993). Two pointsare clear about epithermal deposits: a significant portion ofnear-neutral pH chloride waters is derived from deeply circu-lated meteoric water, and there is evidence in some depositsfor a component of magmatic water, thus a potential source ofsome components, even metals (e.g., Simmons, 1995).

Mineralization affiliated with alkaline rocks

Cripple Creek, Ladolam, Emperor, and Porgera are groupedas a subtype of the deposits associated with quartz ± calcite ±adularia ± illite assemblages but are distinguished becausethey show a number of distinctive features, including associa-tion with alkaline igneous rocks, and the common occurrenceof telluride minerals in their ores (Bonham, 1986; Richards,1995; Jensen and Barton, 2000; Sillitoe, 2002). Although theyare relatively few, these alkaline rock-related deposits havesignificant gold contents and grades, and they display featuressuggesting genetic aspects that differ from most other ep-ithermal deposits formed from near-neutral pH solutions(Table 1). Gold occurs in native form, in electrum, in tel-lurides, and in refractory pyrite, the latter of which can be asignificant component of ores (Carman, 2003; Pals et al.,

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0 10

Au-Ag

Ag-Pb-Zn

Au (Cu)

Au (Te) (alkaline rocks)

wt % NaCl equivalent

20 30

FIG. 8. Fluid inclusion salinities vs. metal contents in epithermal deposits.Gold-silver, gold (Te), and Ag-Pb-Zn deposits are associated with quartz ±calcite ± adularia ± illite gangue, whereas the Au (Cu) deposits are associatedwith quartz + alunite ± pyrophyllite ± dickite gangue.

0

-20

-40

-60

δD

(‰, S

MO

W)

δ18O(‰, SMOW)

-80

-100

-120

-140

-20 -15 -10 -5 0 5 10 15 20

Met

eoric

wat

er

Volcanic vapors

Felsic magmas

Qtz±Calc±Ad±Illite epithermal deposits

Mexican Ag-Pb-Zn epithermal deposits

O-shift water-rock interaction

Lepanto

Ladolam

Qtz-Alun±Pyroph±Dick±Kao epi. deposits

FIG. 9. Stable isotope (δD vs. δ18O) patterns for epithermal deposits (compiled from Arribas, 1995; Simmons, 1995;Cooke and Simmons, 2000; and Albinson et al., 2001). The trend for Lepanto is based on hydrothermal alunite that is a haloto the enargite-bearing ore; the trend indicates condensation of magmatic vapor by local meteoric water (Hedenquist et al.,1998). The trend for Ladolam represents modern geothermal waters and shows mixing between magmatic and local mete-oric water (Carman, 2003). The O shift due to water-rock interaction is based on Taylor (1979).

2003). Adularia is a dominant gangue mineral, probably at theexpense of quartz, which is generally subordinate (Jensen andBarton, 2000), perhaps due to the higher quartz solubilityunder alkaline conditions (Sillitoe, 2002). Fluorite, roscoelite(vanadium-bearing mica), and telluride minerals are com-mon, although not essential accessory minerals, and the oc-currences of magnetite ± hematite, Fe-rich sphalerite, andtetrahedrite-tennantite indicate low- to moderate-sulfidationstates (Jensen and Barton, 2000); lattice and colloform,banded vein textures are rare. Ores extend over unusuallylarge vertical intervals (500–1,000 m) and can be associatedwith telescoping of epithermal and porphyry environments(Jensen and Barton, 2000; Sillitoe, 2002). Hydrothermal al-teration is restricted to areas immediately adjacent to ore,where there is extensive development of propylitic andargillic assemblages. There is also a lack of zoning amongtemperature-sensitive alteration minerals, such as clays. Fluidinclusion studies indicate that ore fluids had salinities of <5 to10 wt percent NaCl equiv, and along with stable isotope data(e.g., see Ladolam trend, Fig. 9), suggest that mineralizingfluids contain a substantial proportion of magmatic water(e.g., Ahmad et al., 1987; Richards, 1995; Simmons, 1995;Carman, 2003; Ronacher et al., 2004).

Mineralization Associated with Quartz + Alunite ±Pyrophyllite ± Dickite ± Kaolinite Assemblages

A second type of epithermal mineralization is distinguishedby its close association with quartz + alunite ± pyrophyllite ±dickite ± kaolinite alteration that forms from hypogene acidicfluids in extinct magmatic hydrothermal systems. This ganguemineral assemblage may be host to Au-Ag ± Cu and Au-Cuores, and rarely Ag-Au-Zn ores (e.g., La Coipa, Chile). Quartzis the dominant gangue mineral, with variable, commonlyhigh, amounts of pyrite and/or marcasite, alunite (Fig. 10),dickite and/or kaolinite, pyrophyllite, diaspore, barite, andalumino-phosphate sulfate (APS) minerals. Ore mineralsoccur as open-space filling and replacement of preexistingminerals. The strongly altered, highly siliceous rock is eitherporous and vuggy or massive and dense (Fig. 10), the latterlargely due to silicification of the residual quartz developeddue to leaching (Steven and Ratté, 1960). Orebodies tend tobe somewhat irregular in shape and strongly influenced byzones of high permeability, which are controlled by structureand/or lithology, comprising vertical structures and pipes andbedded volcanic or sedimentary rocks. Ore also occurs inzones of brecciation. Native gold predominates, whereaselectrum is minor and commonly occurs with copper-arsenicand copper-antimony minerals, such as enargite, its dimorphluzonite, famatinite, tetrahedrite, and tennantite (e.g., Whiteet al., 1995); the first three indicate a high-sulfidation state,the latter two an intermediate-sulfidation state. Parageneticrelationships between these minerals indicate that sulfidationstates fluctuate between high and intermediate but that mostnative gold (not electrum) deposition is associated with inter-mediate-sulfidation states (White et al., 1995; Einaudi et al.,2003), developed after most enargite has deposited; refrac-tory gold commonly occurs in the lattice of sulfide minerals.Other common but typically minor minerals include chal-copyrite, covellite, sphalerite, and tellurides, and locally chal-cocite and bornite.

Ore textures show relatively little variation and are domi-nated by bodies of vuggy and/or massive quartz (Fig. 10). Insome deposits (e.g., El Indio, Chile; Lepanto, Philippines;Chinkuashih, Taiwan), veins and breccias are important hoststo high-grade ore, as are massive occurrences of enargite, lu-zonite, and pyrite (Fig. 10). In other less-common deposits(e.g., Pueblo Viejo, Dominican Republic), vuggy quartz maybe minor or absent, and the dominant alteration is pyrophyl-lite-quartz-pyrite ± alunite, which typically forms deeper inthe hydrothermal system, in some cases close to porphyriticintrusions (Thompson et al., 1986; White, 1991; Mastermanet al., 2002; Gustafson et al., 2004). Supergene oxidation ofsulfides plays a crucial role in liberating gold from refractory

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FIG. 10. Photographs of minerals and textures that commonly occur in ep-ithermal deposits associated with quartz + alunite ± pyrophyllite ± dickite ±kaolinite. A. Oxidized, vuggy residual quartz (Summitville, United States;scale bar = 2 cm). B. Massive enargite (El Indio, Chile; scale bar = 1 cm). C.Coarse tabular alunite on vuggy quartz (Tambo, Chile; scale bar = 1 cm).Note that the vugs in the Summitville sample represent leached feldsparphenocrysts.

A

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sulfides, making mineralization that might otherwise be un-economic (e.g., Pueblo Viejo, Dominican Republic; Nelson,2000) amenable to low-cost processing (e.g., Yanacocha,Peru; Harvey et al., 1999).

Vuggy quartz is a residual product of intense acid alteration,and it is a distinctive feature that reflects the original rock tex-ture and differential leaching of phenocrysts and/or lithicfragments. Its formation predates deposition of copper andgold, which are introduced by a fluid of different composi-tion, illustrating the importance of paleopermeability inpreparation for metal deposition (e.g., White, 1991; Arribas,1995). The vuggy quartz texture in combination with dickiteand/or kaolinite and pyrophyllite indicates that initial fluidscausing alteration and rock dissolution were extremely acid(pH <2 for aluminum to be soluble; Stoffregen, 1987). Thepresence of magmatic hydrothermal alunite indicates that thefluids were relatively oxidized. The vuggy quartz zone flaresupward but may narrow toward the surface where shallowrock units have low permeability, diminishing the alterationeffects of acid-leaching solutions (e.g., Nansatsu; Urashima etal., 1981). In some deposits, such as Cerro de Pasco (Einaudi,1977; Baumgartner et al., 2003) and San Gregorio-Colqui-jirca, Peru (Bendezú et al., 2003), and East Tintic, UnitedStates (Bartos, 1989), the host rocks comprise carbonate unitsand because of their acid-neutralizing capacity, the character-istics of the distal ores differ. In these deposits, vuggy quartzonly occurs in igneous plugs and diatremes that have focusedfluid flow. In adjacent carbonate wall rocks, the base metal-sulfide assemblage is zoned, with a core dominated by high-sulfidation–state Cu sulfide assemblages hosted by siliciczones, grading laterally through intermediate-sulfidation–state Cu-rich assemblages to distal ores dominated by spha-lerite and galena.

The passage of acid solutions leads to broad-scale alterationpatterns that form distinct concentric zones around deposits(Fig. 7). Outward, these zones comprise a silicic core (leachedand silicified rock), alunite, dickite and/or kaolinite or pyro-phyllite, and smectite or illite assemblages (Steven and Ratté,1960). This alteration may extend laterally from 1 to >100 mbut commonly is confined to zones <10 m wide, and theboundary between the central silicic alteration and the outerzones is typically knife sharp (e.g., Summitville, UnitedStates; Nansatsu deposits of Iwato, Akeshi, and Kasuga,Japan; Chinkuashih, Taiwan). Propylitic alteration is wide-spread and surrounds the acid-altered core, but zones of illiteand pyrophyllite can extend well below some deposits (e.g.,El Indio, Chile; Rodalquilar, Spain; Lepanto, Philippines;Yanacocha, Peru). These changes in alteration assemblagesreflect outward and upward neutralization of acid fluidsthrough water-rock interaction and cooling (Steven andRatté, 1960). In some districts and deposits (e.g., El Indio-Pascua belt, Chile-Argentina; Bissig et al., 2002, Deyell et al.,2004; Puren prospect near La Coipa, Chile; Sillitoe, 2004),the shallow level is preserved and represented by steam-heated advanced argillic alteration that marks the position ofthe paleowater table (Sillitoe, 1993b, 1999).

Fluid inclusion data

Because of the corrosive nature of acid solutions and thecharacter of hydrothermal minerals produced during the

early leaching (i.e., alunite, kaolinite, pyrophyllite), and thelack of quartz associated with the ore stage, fluid inclusionsrepresenting ore-forming conditions for these deposits arerare. Based on the infrared microscopy of enargite, discussedbelow, mineralization appears to be associated with dilute tomoderately saline solutions (<5 to ~10 wt % NaCl equiv; Fig.8) at temperatures between 180° and 320°C (Deen et al.,1994; Arribas, 1995; Mancano and Campbell, 1995; Deyell etal., 2004), with hotter and more saline fluid inclusions occur-ring sporadically, especially in deep rocks beneath ore zones(Arribas et al., 1995a, b; Bethke et al., 2005). Hydrostatic boil-ing point for depth gradients appear to prevail, suggestingthat ores formed between ~100 and ~1,000 m below the pa-leowater table (Arribas, 1995; Sillitoe, 1999). There are datafor late, coarsely crystalline quartz and secondary inclusionsin quartz phenocrysts; coexisting liquid- and vapor-rich inclu-sions are commonly observed, indicating boiling conditions(e.g., Arribas et al., 1995b; Ruggieri et al., 1997; Wang et al.,1999).

Fluid inclusion studies on enargite provide much of the di-rect evidence of the salinity of ore-depositing solutions. De-terminations for Julcani (Deen et al., 1994) range 8.1 to 19.2wt percent NaCl equiv (Th 200°–300°C) and for Lepanto(Mancano and Campbell, 1995) range from 0.5 to 4.5 wt per-cent NaCl equiv (Th 190°–290°C), consistent with more re-cent data reported for deposits in Bulgaria and Serbia (seeSillitoe and Hedenquist, 2003; Heinrich, 2005). The salinity-homogenization temperature data from Lepanto representone of the most detailed data sets available; the data plot in abroad linear trend indicating enargite deposited during cool-ing of the parent fluid (~300°C, ~5 wt % NaCl equiv), causedby mixing and dilution with shallow ground water during lat-eral flow along the strike of the deposit (Mancano and Camp-bell, 1995).

Stable isotope data

Stable isotope measurements of δD, δ18O, and δ34S havebeen determined for OH-bearing phases (alunite, illite, andkaolinite) and coexisting sulfides (Rye et al., 1992; Arribas,1995; Hedenquist et al., 1998; Deyell et al., 2004; Bethke etal., 2005; Fifarek and Rye, 2005). The δD versus the δ18Odata for the advanced argillic alteration cluster tightly nearthe composition of magmatic vapor (Giggenbach, 1992b) andshow mixing trends with meteoric water (Fig. 9). The signifi-cant involvement of a magmatic component is consistent withore formation at relatively shallow levels above crystallizingand degassing stocks. In some cases, the mixing trends sup-port the fluid inclusion data, indicating that dilution and theconcomitant cooling influence mineral deposition (e.g.,Hedenquist et al., 1998). δ34S measurements on sulfate-sul-fide pairs indicate equilibration temperatures of 200° to>350°C and a magmatic source of sulfur (Rye et al., 1992; Ar-ribas, 1995; Hedenquist et al., 1998; Deyell et al., 2004;Bethke et al., 2005; Rye, 2005).

Association with Intrusion-Centered DepositsEpithermal deposits are the shallow and relatively distal

part of a continuum of ore-forming environments that form inhigh-temperature hydrothermal systems, of which the mostproximal, deepest, and hottest part is represented by intrusion-

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centered ore mineralization such as porphyry Cu (Mo-Au),carbonate replacement deposits, and skarns (e.g., Meinert etal., 2005; Seedorff et al., 2005). High- to intermediate-sulfi-dation–state sulfides and ores that are associated with quartz+ alunite ± pyrophyllite ± dickite ± kaolinite gangue typicallyshow the strongest links to these more deeply formed stylesof mineralization, as is indicated by their common occurrencein proximal volcanic settings, their overlapping wall-rock al-teration and ore mineral assemblages, and overlapping fluidinclusion characteristics (Sillitoe, 1999; Einaudi et al., 2003;Sillitoe and Hedenquist, 2003; Heinrich, 2005). There aredistricts (e.g., Mankayan district, Philippines; Cajamarca re-gion, northern Peru; Maricunga belt and Collahuasi district,Chile; Ag-Pb-Zn belt of northern Mexico) where intrusion-hosted mineralization and epithermal mineralization occurclose to one another (Megaw et al., 1988; Arribas et al., 1995a;Hedenquist et al., 1998; Sillitoe et al., 1998; Muntean andEinaudi, 2001; Gustafson et al., 2004; Masterman et al., 2004).In some cases, the adjacent occurrences of mineralizationhave nearly the same age, implying a close genetic link be-tween hydrothermal activity, igneous intrusions, and ore min-eralization, such as in the Mankayan district, Philippines (Ar-ribas et al., 1995a; Hedenquist et al., 1998; Claveria, 2001;Sajona et al., 2002). Rarely, epithermal mineralization is foundclose to deep batholiths (Kesler et al., 2004). This continuumof mineralization styles, however, should not be taken as indi-cation that an observed close spatial correlation between shal-low and deep formed ores is common, despite the likelihoodthat igneous intrusions underlie (probably at considerabledepth) a significant portion of epithermal deposits (Fig. 1).

Case StudiesWe review five examples to illustrate the variation between

epithermal deposits, highlighting the geologic context, min-eral zoning, and the range of metal occurrences and patternsat different scales. We start with the very large Ag-rich (Au)Pachuca vein deposit in Mexico followed by the intermediate-size Chinkuashih Au (Cu) deposit of Taiwan to emphasize thedeposit-scale controls on ores of the two different types of ep-ithermal mineralization. This is followed by a description ofthe giant Yanacocha Au (Cu) deposit in northern Peru, simi-lar in origin to Chinkuashih but with distinctly different litho-logic controls. We review the Hauraki goldfields of the Coro-mandel peninsula, North Island, New Zealand, and theepithermal belts of the western United States and northernMexico, to show provincial characteristics, clustering of de-posits, and variations in metal contents, at the regional tometallogenic belt scale.

Pachuca-Real del Monte, Mexico

Pachuca-Real del Monte has the largest silver productionfrom an ore deposit. It covers an area of about 100 km2 andhas produced ~45,000 t Ag and 220 t Au (Fig. 11). Similar tomany of the early discoveries in Latin America, it was minedby the Spanish from about 1550, although it was likelyworked in pre-Hispanic time. The deposit was exploited byunderground mining, but these activities are now greatly re-duced and small in scale. The district is divided into two min-ing areas, Pachuca to the west and Real del Monte to theeast. Excellent descriptions of the geology, alteration, and

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mineralization are found in Geyne et al. (1963) and Dreier(1976, 1982).

The deposit is entirely hosted by a thick sequence (up to1,000 m) of calc-alkaline volcanic and hypabyssal rocks, rang-ing from basaltic andesite to rhyolite. These rocks are region-ally altered to a propylitic assemblage of quartz, epidote, chlo-rite, adularia, albite, calcite, and pyrite, but near the veins thehost-rock alteration forms narrow envelopes a few meterswide comprising quartz, adularia, and pyrite. Orebodies arecontained in a series of east-west–, northwest-southeast–, andnorth-south–trending, fault-hosted veins (Fig. 11). Individualveins range from 0.5 to ~5 m in width, although there aresheeted zones comprising closely spaced veins filling frac-tures that range up to 35 m in width. Major faults have nor-mal offsets of up to 350 m, but many of the faults hostingveins show displacements of only a few meters. These faultswere active both before and during mineralization, and al-though some are recognizable on the surface, many are diffi-cult to trace and poorly exposed.

Ore is localized in dilatant, highly fractured zones, formingirregular pods and lenses separated by low-grade and barrenzones where the fault zone narrows or is filled with gouge.Productive ore zones occur where there are changes in veinstrike and dip. Throughout the district, ore is constrained toan ~800-m vertical interval (2,800–2,000 m asl), making up a“mineralized horizon,” although the vertical extents of oresvary widely between veins (Fig. 12). In many veins, ore ter-minates upward below the surface into zones of intense clayalteration comprising illite, calcite, chlorite, and pyrite withdiscontinuous veinlets of quartz and calcite (Fig. 13). Down-ward, orebodies terminate where veins narrow to less thanminimum mining width, where they pinch out, or where sil-ver values decrease below the cutoff grade; deep drillingshows that some veins extend several hundred meters be-neath the ore zone. Representative ore mined in the mid-1930s contained 410 g/t Ag, 2.3 g/t Au, 0.5 wt percent Zn,0.05 wt percent Pb, and 0.01 wt percent Cu (Geyne et al.,1963).

Ore was deposited during several episodes of vein filling,forming coarse-scale crustiform banding of fine- to medium-grained quartz; colloform banding and lattice textures havenot been reported. Quartz, with minor amounts of adularia,albite, calcite, pyrite, chalcopyrite, galena, and sphalerite arethe common vein minerals, with silver occurring mainly inacanthite. This mineralogy reflects the reduced near-neutralpH and intermediate-sulfidation state of the ore-forming so-lution. Mn-bearing minerals, rhodonite and bustamite, arepresent in trace amounts in many veins but occur in signifi-cant abundance where there is rich silver ore. Sulfides gener-ally constitute between <1 and about 12 vol percent of thevein fill. Metal zoning has not been reported except that basemetals are generally more abundant at depth; the combinedbase metal content never exceeded 4 wt percent (Geyne etal., 1963). Fluid inclusion homogenization temperaturesrange from 210° to 305ºC, and ice-melting temperatures in-dicate hydrothermal solutions had salinities of 0 to 6 wt per-cent NaCl equiv (Dreier, 1976). Postore deformation is re-stricted to strike-slip displacements, which occurred alongthe east- and northwest-trending veins, offsetting north-trending veins by <1 to >100 m.

As is characteristic of many Mexican epithermal deposits,the veins are very long. The east-west–trending Vizcaína vein(Fig. 11) is mineralized for an 8-km strike length and ishosted by a major fault that crosses the district. Numerousdikes, <5 to 100 m wide, predate mineralization and followfaults parallel to the east-trending veins, indicating that mag-matism and hydrothermal activity shared the same stressregime. Most dikes are fine grained and are intrusive equiva-lents of the overlying volcanic rocks; a small intrusion consist-ing of coarse-grained quartz porphyry occurs locally in thesoutheastern part of the district (Fig. 11). Precious metalmineralization formed at about 20.3 Ma (McKee et al., 1992),during the last phase of ore deposition within the Mexican sil-ver belt and during the last phase of felsic volcanism withinthe Sierra Madre Occidental ignimbrite province (Camprubiet al., 2003). Despite being part of a much larger metallogenicbelt, described below, these age relationships suggest that thePachuca-Real del Monte formed in isolation, ~500 km fromSan Martín de Bolaños, which is the nearest known deposit ofsimilar age (Camprubi et al., 2003).

Chinkuashih, Taiwan

Chinkuashih is a Pliocene-age deposit located near thenortheast tip of Taiwan at the junction of the Luzon andRyukyu arcs. Epithermal mineralization was discovered in1894 and was exploited until the mid 1980s. Historic produc-tion exceeds 90 t gold, with an unknown quantity of copper.Gold-copper mineralization occurs in a zone of intensesiliceous alteration that forms some of the highest peaks, oneof which (since removed by mining) gave the deposit its name(Chinkuashih means golden melon stone). Until the mineclosed in 1987, ore was treated by flotation of all sulfide min-erals to produce a pyrite concentrate containing approximately1 wt percent Cu, mostly in the form of enargite, and bycyanide treatment of the tailings. Huang (1955, 1963), Folins-bee et al. (1972), Huang and Meyer (1982), Tan et al. (1987),and Tan (1991) described the characteristics of the deposit.

The host rocks are augite-biotite-hornblende phyric dacitebodies intruded into a sequence of moderately steep west-dipping shallow marine to continental sedimentary rocks con-sisting of calcareous sandstones, siltstones, conglomerates,mudstones, sedimentary breccias, and minor coal beds (Fig.14). These rock types host more than 80 individual orebodies.The largest, Penshan, was responsible for the bulk of produc-tion. It occupies a north-south zone about 2 km long, dipping70° E, which cuts across an intrusive dacite dome and extendsinto the adjacent sedimentary rocks.

Mineralization is confined to zones of intense silicic alter-ation (vuggy quartz, with or without alunite, and massivequartz), typically with abundant pyrite (Fig. 14). North-southfaults, breccia pipes and favorable stratigraphic horizons con-trol ore occurrences in the Penshan area. Average gradeswere 0.55 wt percent Cu and 2.7 g /t Au. The most favorablehost rocks are dacite and calcareous sandstone. In the dacite,ore is hosted by vuggy quartz formed by acid leaching butwith strong structural controls. In the calcareous sandstone,pyrite completely replaces the calcareous matrix, and orezones that are narrow in other rock types widen sharply.There is a great diversity of structural controls on the manyindividual orebodies at Chinkuashih, including bedding

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planes, faults, joints, and the margins of intrusions. Gold isfound in veins and pipes filled with massive enargite andpyrite and disseminated in siliceous bodies; bonanza zonesoccur at shallow depths in breccia pipes. Overall, gold

contents are higher at shallow levels, whereas copper con-tents are higher at deeper levels.

The Penshan main vein is the principal orebody, with oreextending over a 650-m vertical interval in strong association

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Pachuca - Real del Monte district - vertical extent of mineralized veins

FIG. 12. Vertical distribution of ores in veins of the Pachuca-Real del Monte mining district (redrawn from Geyne et al.,1963). Their different levels of formation are inferred to reflect shifts in the pressure and temperature of hydrothermal con-ditions conducive to metal precipitation (Geyne et al., 1963); there is no indication that they are a product of postore faultdisplacements.

with zones of massive and vuggy quartz. Locally, these zonesattain widths of 100 m. Enargite forms striated prismatic crys-tals up to 2 cm long, intergrown with pyrite (pyritohedral andoctahedral forms) and barite. Gold occurs mainly in its nativeform with minor electrum and trace amounts of tellurides;auriferous pyrite contains <1 to 100 ppm Au (Tan, 1991).Massive sulfide ores contain pyrite, enargite, luzonite, andminor chalcopyrite, covellite, tetrahedrite, and chalcocite.Quartz and alunite are common gangue minerals. Sphaleriteand galena are rare; cinnabar and native mercury occur nearthe surface at Bamboo and Changjen.

Zones of silicic alteration (i.e., vuggy quartz, locally withsubsequent quartz flooding) are surrounded by alterationdominated by kaolinite and dickite with minor nacrite, alu-nite, quartz, and pyrophyllite, which give way outward to al-teration consisting of kaolinite, montmorillonite, and mixed-layer clay (Chen, 1971). These clay-dominated assemblagesoccupy relatively narrow zones no more than a few meterswide (Fig. 14) and grade sharply into widespread chlorite-car-bonate alteration that, at distances of several hundred meters,grades into fresh dacite (Huang, 1955). This alteration se-quence reflects the restricted and focused nature of fluid flowand intense acid leaching, which enhanced the permeability.Fluid inclusion data on barite intergrown with enargite,

hydrothermal quartz, and igneous quartz (secondary inclu-sions) indicate homogenization temperatures of 190° to290°C that are consistent with equilibration temperaturescalculated from δ34S data for barite-sulfide (enargite, pyrite)pairs (Folinsbee et al., 1972; Wang et al., 1999). The bulk ofthe ice melting temperatures indicate salinities of 0.5 to 5.0wt percent NaCl equiv, but there are some halite-bearing liq-uid-rich inclusions, with vapor-rich inclusions, that homoge-nize from 300 to >370°C (Wang et al., 1999).

Quartz-carbonate veins at Chiufen, with historic produc-tion of over 30 t Au, occur about 1.5 km from the copper-goldepithermal deposits of Chinkuashih, and their proximity toeach other suggests they may have formed about the sametime. The veins contain electrum, pyrite, chalcopyrite, stib-nite, sphalerite, galena, rhodochrosite, and barite, typical ofintermediate-sulfidation veins, and represent formation fromreduced, near-neutral pH solutions. Such side-by-side occur-rences of contrasting alteration styles (acid vs near-neutral pHgangue assemblages) are relatively common, albeit with onetypically hosting a minor metal resource compared to theother, as occurs in Japan, Peru, and western Great Basin (Sil-litoe, 1999; John, 2001). Chinkuashih-Chiufen is an exceptionwhere both epithermal environments contain mined ore de-posits, as is also the case with Lepanto and Victoria (Claveria,2001).

Yanococha, Peru

Yanococha, located in northern Peru (3,300–4,200 m asl), isone of the premier epithermal districts in the world, contain-ing a total resource that exceeds 1,300 t of gold (Fig. 15). Al-though primitive mining for mercury dates back to 300 AD, itwas not until the early 1980s that systematic gold explorationwas initiated and geochemical anomalies were identified bydrainage, soil, and rock-chip sampling (Sillitoe, 1995a; Harveyet al., 1999). There are at least eight distinct epithermal golddeposits in Miocene volcanic rocks known in the district, withseveral mined from open pits (Harvey et al., 1999; Bell et al.,2004). In addition, there is a related gold deposit (La Quinua)that is hosted entirely by young unconsolidated basin-fillgravels derived from glacial-fluvial erosion of bedrock con-taining epithermal mineralization (Mallette et al., 2004). Ep-ithermal gold ores are disseminated, oxidized by supergeneweathering, and have average grades of 0.8 to 1.6 g/t. Recentsummaries describing the deposit and its geologic character-istics are found in Turner (1997), Harvey et al. (1999), Bell etal. (2004), and Longo (2005).

The Miocene Yanococha Volcanic Complex that hosts thedeposits covers an area approximately 20 by 25 km and wasformed between about 14.5 and 8.4 Ma (Longo, 2005). Thecomplex consists of numerous domes, dikes, and diatremesthat intrude a sequence of pyroclastic deposits and lava flowsup to several hundred meters thick. Rocks range from an-desite to dacite and contain phenocrysts of feldspar, pyrox-ene, amphibole, and quartz. Numerous breccias fill dikes andpipes and many contain round to subangular clasts and fine-grained matrix suggesting explosive emplacement fromphreatic and phreatomagmatic eruptions (Bell et al., 2004). Afew of the diatremes contain distinct feldspar porphyry claststhat indicate the presence of porphyry copper mineralizationat depth (Harvey et al., 1999; Gustafson et al., 2004). These

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100 m

argillic

K-spar

propylitic

vein

FIG. 13. Schematic cross section of the Pachuca-Real del Monte miningdistrict, showing the relationship between argillic, K-feldspar, and propyliticalteration patterns and blind veins (redrawn from Dreier, 1982).

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2 km0

92 t Au182 t Ag

Chinkuashih

Au veinDacite

Cu-Au ore

Changjen

Penshan

ChiufenPine

Tortoise

Buffalo

200 m0

300

400

500

200

100

W E

ShaleCalcareous sandstone

vuggyquartz

clay alteredselvage

clay alteredselvage

BA

C

Bamboo

FIG. 14. Chinkuashih epithermal deposit, Taiwan. A. Sketch plan of main geologic features showing the locations of ore-bodies and dacite intrusions (modified from Tan, 1991). B. Sketch cross section showing structural and lithological controlon mineralization within the Penshan ores (the eastward dip on the strata shown is a local feature that differs from the west-dipping regional trend; from Huang, 1963). C. Vertical zone of gold mineralization cutting shallow-dipping interbeddedsandstone and shale adjacent to the Penshan dacite. The central (dark-colored) acid leached siliceous zone grades 3.8 g/t Au.The pale-colored zones adjacent to it are barren clay altered zones, which grade sharply into unaltered wall rock over about1 to 2 m.

igneous rocks are associated with regional early to middleMiocene magmatism that developed in northern Peru (Nobleand McKee, 1999). Their remnants underlie the YanacochaVolcanic Complex and comprise andesitic lavas, lahars, anddebris flows. An unconformity separates them from the un-derlying, thick and strongly folded Cretaceous sedimentarysequence of quartzite, argillite, and limestone.

Progressive deformation, extending from the Paleocene tolate Miocene, has affected all the rocks in the region (Bena-vides-Cáceres, 1999). Faults and fractures within the district

strike northeast-southwest, northwest-southeast, and east-west, and were active multiple times. The northwest-south-east structures show evidence of both transcurrent movementdue to compression and normal movement due to extension,whereas the other main trends relate only to open fracturesand normal faults due to extension. Mineralized zones, brec-cias, and shallow intrusions are aligned in a northeast-trend-ing belt that is part of the larger trans-Andean Chicama-Yana-cocha corridor, which is 30 to 40 km wide and extends for 200km from the Cajamarca region to the Pacific Ocean (Fig. 15;

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Cajamarca

Yanacocha

Tertiary volcanic rocks

Intrusions

Paleozoic-Mesozoicsedimentary rocks

Andean parallelstructures

Trans-Andeanstructures

Epithermal

Porphyry Cu-Au-Mo

Polymetallic Ag

4 km

Silicification (massive, vuggy, granular, opaline)

Silica alunite

Silica clay

Clay

Propylitic/fresh rock

Ore body

A

B

NLa Quinua

Fault

FIG. 15. A. Map of the Cajamarca province showing the locations of Yanacocha and surrounding epithermal and porphyryCu-Au-Mo deposits, which are localized at the intersection of northwest-trending Andean and northeast-trending Trans-An-dean structures (modified from Gustafson et al., 2004). B. Map of the Yanacocha district and the outline of orebodies withhalos of hydrothermal alteration (from Bell et al., 2004); La Quinua is an alluvial deposit containing >10 Moz gold (Malleteet al., 2004).

Quiroz, 1997). Its intersection with the Andean fold belt mayhave played a role in focusing magmatism and hydrothermalactivity in the Yanacocha district (Turner, 1997). Locally, east-west– and north-south–trending fracture zones of interpretedextensional origin host high-grade gold ores (Bell et al., 2004).

Hydrothermal alteration is pervasive and widespread in thedistrict. Gold mineralization is associated with tabular, sub-horizontal masses of intense silicic alteration due to acidleaching, largely of welded pyroclastic units. Silicic alterationincludes granular quartz, vuggy quartz, and massive quartz,which is zoned vertically and horizontally and attains thick-nesses of many tens to several hundreds of meters (Harvey etal., 1999). Hypogene alunite is zoned outward and downwardfrom these siliceous masses, extending several tens to severalhundred meters from the silicic zone, replacing groundmassand phenocrysts as well as forming fine crystals growing intoopen space. Pyrophyllite occurs locally intergrown with alu-nite and grades outward from pyrophyllite-alunite to pyro-phyllite-alunite-kaolinite to pyrophyllite-kaolinite. Kaoliniteand smectite clays dominate the distal zones of acid alterationbut give way to regional propylitic alteration containing chlo-rite, smectite, quartz, illite, pyrite, and calcite.

Gold was introduced as disseminations, fracture and vug fill-ings, or in breccias during several stages of hydrothermal ac-tivity, spanning ~5.4 m.y. from 13.6 to 8.2 Ma (Longo, 2005)and postdating the phases of intense leaching and silicic alter-ation. Submicron- to micron-sized particles of native goldwere deposited with fine-grained pyrite and minor enargiteand covellite. Later gold deposited with pyrite, enargite, cov-ellite, alunite, and barite. Patchy occurrences of veins filledwith rhodochrosite, dolomite, and base metal sulfides repre-sent the last hydrothermal event and a shift to near-neutral pHconditions (Bell et al., 2004). A significant portion of ore wassubject to deep supergene oxidation (>300-m depth) duringwhich iron oxide minerals formed at the expense of sulfideminerals (Harvey et al., 1999). Oxidation, coupled with thefact that gold postdates most silica deposition, thus precludingsilica encapsulation, means that metallurgical processing ofoxide ores is relatively straightforward (Bell et al., 2004).

Multiple centers of porphyry copper-gold mineralizationunderlie epithermal orebodies in the Yanacocha district (Fig.15). Drilling results from Kupfertal show that near-surfacepyrophyllite-alunite-kaolinite alteration grades downwardinto weak sericite-chlorite alteration which overprints weakpotassic alteration at depth (Gustafson et al., 2004), associ-ated with porphyry-style mineralization of 0.3 wt percent Cuand 0.3 g/t Au.

Hauraki goldfields, New Zealand

The Hauraki goldfields in the Coromandel peninsula oc-cupy an area of about 200 × 35 km and containing a denseclustering of epithermal deposits and several exposures ofweakly developed porphyry Cu-style mineralization (Fig. 16).More than 50 deposits were mined between 1852 and 1952,during which time approximately 1,400 t of bullion (Au/Ag =1/3) was produced. Christie and Brathwaite (1986), Brath-waite et al. (1989), Brathwaite and Christie (1996), Brath-waite et al. (2001a, b), and Brathwaite and Faure (2002)described the regional setting and characteristics of theepithermal deposits.

Precious metal mineralization is hosted by a thick sequence(up to 2,000 m) of Miocene to Pliocene volcanic rocks thatdip at a low angle to the east. Andesite flows and brecciasdominate the lower part of the sequence, whereas rhyoliticlavas and pyroclastic deposits dominate the upper part. Thevolcanic rocks rest unconformably either on a thin sequenceof Oligocene-early Miocene marine sedimentary rocks or onbasement composed of Mesozoic graywacke and argillite ex-posed locally in the northern part of the peninsula. Faults dipsteeply and strike northwest and northeast. Volcanism and hy-drothermal activity that produced widespread alteration halosand epithermal mineralization were both products of subduc-tion-related magmatism that has since migrated southeast tothe now active Taupo Volcanic Zone. Preservation of the pa-leosurface, which is marked by silica sinter occurrences, is re-stricted largely to the younger exposures on the east side ofthe province.

By far the most important deposit is Martha Hill (see Figs.3, 5), which has produced more than 260 t Au and more than1400 t Ag. Other major producers include: Thames (45 t Au,27 t Ag), Karangahake (29 t Au, 97 t Ag), and Golden Cross(23 t Au, 73 t Ag). All of the ore occurs in veins or veinlets,many of which have a northerly to northeasterly trend. Min-eralization in the Hauraki goldfields generally is associatedwith dominantly normal faults, although in very few casesslickensides indicate strike-slip movements. Veins branch up-ward and also sideways at their extremities, where some ter-minate against crosscutting faults. Most veins are 0.2 to 1.5km long, and range from 0.3 to 5 m in width. Veins extend toconsiderable depth, but because ore usually bottomed outabove their termination, the characteristics at the deepestlevel are poorly known. Ore generally is restricted to verticalintervals of <300 m, except at Martha Hill and Karangahakewhere it extends over intervals of 575 and 700 m, respectively.

Electrum and acanthite are the main gold and silver min-erals. Sulfide minerals comprising pyrite, sphalerite, galena,and chalcopyrite make up <2 to 10 vol percent (up to 30 vol% very locally) of the vein fill. The Tui mine, which was ex-ploited until the early 1970s, is a base metal deposit with sul-fide-rich veins, from which 163,000 t of ore was producedthat contained 4.5 wt percent Zn, 2.7 wt percent Pb, 0.2 wtpercent Cu, 15.2 g /t Ag, and 0.36 g/t Au. Base metal sulfide-rich mineralization also occurs in the deeper parts of bothMartha Hill and Karangahake. In the 1950s and 1960s, Pb-Zn± Au-Ag were prospected in the Sylvia and Monowai de-posits. Other trace minerals include selenides, tellurides, sul-fosalts, stibnite, molybdenite, marcasite, arsenopyrite, andcinnabar, but these occur in only a few deposits. Quartz, cal-cite, and adularia are the main gangue minerals, with sporadicoccurrences of rhodochrosite, inesite (Mn zeolite), siderite,barite, kaolinite, and anhydrite.

Martha Hill comprises a branching system of subparallel,steeply dipping quartz-calcite-sulfide-adularia veins that ex-tend about 1.5 km along strike (Fig. 3). The Martha vein islargest and up to 35 m wide. Mining began in 1882 followingdiscovery of surface exposures of veins, but most of the min-eralization was concealed by postore rhyolitic ignimbrite andunderlying lake sediments that fill a paleodepression runningdown the axis of the vein system (Fig. 4). Most of the ore wasextracted by underground mining (1883–1952), but since

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1988, vein material left behind and stope fill have been ex-ploited by open-pit mining. Fine-grained electrum-sulfidegold-silver ore contains 64 g/t Au, 400 g/t Ag, 0.47 wt percentZn, 0.41 wt percent Pb, and 0.15 wt percent Cu based on arepresentative assay of drill core over 1 m, in comparison todeep coarse-grained sulfide ore, which contains 0.1 to 2 g/tAu, 5 to 70 g/t Ag, 0.3 to 4 wt percent Zn, 0.05 to 2 wt per-cent Pb, and 0.01 to 0.2 wt percent Cu based on assay of drillcore over widths up to 8 m (Brathwaite and Faure, 2002). Theenclosing andesitic host rocks are extensively altered over an

area of >10 km2, and in the vicinity of veins they are alteredto adularia, illite, quartz, calcite, chlorite, and pyrite, withminor albite and epidote. The nearby Favona, Karangahake,and Golden Cross deposits are similar in terms of mineraliza-tion style, alteration and gangue mineralogy, and silver/goldratio, although base metal sulfides are lacking at GoldenCross.

The Thames deposit is notable for its bonanza ore shoots,which coincide with fault-vein intersections. Epithermal min-eralization there is closely associated with subeconomic

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25 km

N

Hauraki R

ift

Alluvial sediments

Late Miocene-Pliocenerhyolite

Early Miocene-Plioceneandesite and dacite

Early Miocenediorite intrusion

Jurassic greywackebasement

Fault

Epithermal Au-Ag

Porphyry Cu

Epithermal Au-Ag (selenium)

Epithermal Au-Ag (tellurides)

Martha Hill-Favona

Thames

Tui

Karangahake

Golden Cross

Coromandel

Auckland

Wellington

Waihi

50 km

North Island

Main map

Epithermal Pb-Zn (Ag-Cu)

peninsula

FIG. 16. Map of epithermal and porphyry deposits in the Hauraki goldfields and the Coromandel peninsula, North Is-land, New Zealand. The steep, rugged topography is covered by dense vegetation and the rocks are deeply weathered, in-cised by streams and rivers. All the epithermal deposits are associated with quartz ± calcite ± adularia ± illite assemblages(modified from Christie and Brathwaite, 1986; Brathwaite et al., 1989; Brathwaite and Faure, 2002). Epithermal Au-Agdeposits with occurrences of tellurides or selenium-rich sulfides are distinguished (A.B. Christie and M.P. Simpson, pers.commun., 2005)

porphyry Cu mineralization and advanced argillic alteration,indicating telescoping of hydrothermal environments (Brath-waite et al., 2001b). Electrum and pyrargyrite are the maingold-silver minerals along with minor amounts of gold-silvertellurides. Pyrite, sphalerite, stibnite, arsenopyrite, chalcopy-rite, marcasite, enargite, and cinnabar are present in veinsconsisting mostly of quartz and barite.

Within the Hauraki goldfields, there are a few sporadic oc-currences of magmatic hydrothermal advanced argillic alter-ation but none so far have proven to contain significant pre-cious or base metal mineralization. All of the known oresformed from reduced, near-neutral pH solutions of mainly in-termediate-sulfidation state (Christie and Brathwaite, 1986).

K-Ar and 40Ar-39Ar dates for hydrothermal minerals relatedto ore deposition show that the Martha Hill, Favona, GoldenCross, Karangahake, and Tui deposits formed within a shortperiod between 6 and 7 Ma (Brathwaite and Christie, 1996;Mauk and Hall, 2003, 2004). This suggests that several adja-cent ore-forming hydrothermal systems were active contem-poraneously. By contrast, the Thames deposit formed around11.6 to 10 Ma (Brathwaite et al., 2001b).

The Hauraki goldfields, with their numerous occurrencesof epithermal mineralization, highlight the regional and localgeologic controls on ore mineralization. The presence of Au-Ag deposits with telluride and selenide minerals (e.g.,Thames) or of an epithermal Pb-Zn deposit (Tui mine),among numerous other Au-Ag deposits poor in tellurides, se-lenides, and base metals, despite relatively uniform regionalstratigraphy, suggests that deep-seated factors, rather thanlocal country rocks, dictated the metal inventories of ore-forming systems (cf. John, 2001, for a similar distinction inthe Great Basin in western United States). At the same time,the prevalence of vein-hosted ores and mineralogic evidenceof boiling emphasizes the prominence of shallow hydrology inlocalizing orebodies.

Metallogenic belts of western United States-northern Mexico

This large region of southwest North America hosts nu-merous epithermal deposits, comprising a huge endowmentof precious metals concentrated mainly in the Great Basin ofNevada and eastern California and in northern Mexico(Buchanan, 1981). We examine only four provinces to showhow the continuity of metal assemblages forms distinct met-allogenic belts extending hundreds of kilometers. Figure 17shows the distribution of major mid to late Tertiary ep-ithermal deposits and their metal assemblages, allowingcomparison to equivalent-scale tectonomagmatic features,including the distribution of silicic volcanism and Basin andRange extensional tectonics that overlapped the timing ofore formation.

John (2001) described the Miocene to Pliocene epithermaldeposits of the Great Basin and showed how sulfide mineral-ogy of ore correlates with the composition of coeval magmasand plate-tectonic setting. He distinguished two igneous as-semblages and two corresponding metallogenic belts, termedwestern andesite (22–4 Ma) and bimodal basalt-rhyolite (17Ma to Holocene). Western andesite-related deposits includeComstock Lode, Bodie, and Tonopah, with quartz ± calcite ±adularia ± illite gangue assemblages, and Goldfield and Par-adise Peak with quartz + alunite ± pyrophyllite ± dickite ±

kaolinite gangue assemblages. These deposits formed in asso-ciation with the K-rich calc-alkaline continental arc magma-tism, comprising stratovolcanoes, flow domes, and intrusions,in a narrow belt of transtensional and strike-slip faultingknown as the Walker Lane. Amphibole and biotite are char-acteristic hydrous phenocryst phases in the igneous rocks thatoccur with plagioclase, clinopyroxene, orthopyroxene, andFe-Ti oxides (titanomagnetite > ilmenite).

Bimodal basalt-rhyolite–related deposits include Midas,Sleeper, Mule Canyon and the National district, which allhave quartz ± calcite ± adularia ± illite gangue assemblages.They formed during Basin and Range extension and magma-tism in association with mafic dike swarms and lavas and fel-sic flow dome complexes. Basalts contain plagioclase, olivine,and Fe-Ti oxide (ilmenite > titanomagnetite), whereas coge-netic rhyolites locally contain quartz and Fe-rich olivine. Hy-drous mineral phenocrysts are absent in most rocks of allcompositions.

The phenocryst mineralogy associated with the two igneoussuites indicates that the western andesite magmas were moreoxidized and water rich than the bimodal basalt-rhyolite mag-mas (John, 2001). There are also distinctions in mineralogyand chemical signatures of the associated epithermal depositsthat correlate with magma composition. In the western an-desite assemblage, most Au-Ag (Cu) deposits with quartz +alunite ± pyrophyllite ± dickite ± kaolinite gangue containpyrite and/or marcasite, native gold, enargite-luzonite, spha-lerite, covellite, ± chalcopyrite, galena, tetrahedrite-tennan-tite, bismuthinite, stibnite, and gold tellurides, indicating anintermediate- to high-sulfidation state. They generally havesilver/gold ratios of 0.2 to 2 and chemical signatures of Au,Ag, As, Sb, Pb, Cu ± Bi, Hg, Mo, Sn, Te, and Zn. In this samemagmatic assemblage, there also are Au-Ag (Cu-Pb-Zn) de-posits with quartz ± calcite ± adularia ± illite gangue. Thesedeposits contain pyrite, electrum, acanthite, silver sulfosalts,tetrahedrite-tennantite, galena, Fe-poor sphalerite, and chal-copyrite of intermediate-sulfidation state affinity. They gener-ally have silver/gold ratios of 10 to 100 and chemical signa-tures of Au, Ag, Ba, Mn, ± Cu, Pb, Se, and Zn. By contrast,the bimodal basalt-rhyolite–related deposits are relatively sul-fide poor where hosted by rhyolites but relatively sulfide richwhere hosted by basalts. They contain pyrite, marcasite, ar-senopyrite, electrum, acanthite, silver selenides, and minorchalcopyrite, stibnite, and Fe-rich sphalerite of low-sulfida-tion–state affinity and chemical signatures of Au, Ag, As, Sb,Se, Hg, ± Mo, Tl, and W. Silver/gold ratios are generally ≤10(commonly near 1). These mineralogical and geochemicalcharacteristics correlate closely with magma composition;water-poor and reduced mafic magmas correspond to the de-velopment of dominantly low-sulfidation–state ore mineralassemblages, whereas water-rich and oxidized intermediatemagmas correspond to the development of intermediate- (tohigh-) sulfidation–state assemblages.

The work on the Great Basin deposits described above hasno known counterpart in northern Mexico or anywhere elsein the world. Nevertheless, Clark et al. (1982) and Staude andBarton (2001) show, in a more general way, that metallogenythroughout the Mexican region correlates in space and timeto the magmatic evolution associated with plate convergenceduring the mid to late Tertiary. Isotopic dating and fluid

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S

CLC

MO

MS

TA

GO PA

FO

RM

BU

MCI

MN

SL

GFBD

CL

OM

Walker Lane

FW

TU

ML

TAUPP

SB RAZA

Basin and Range province

Middle-Tertiary ignimbrite

Mexican volcanic belt

Geology

130°

20°

20°

30°

40°

40°

120°

120°

110°

110°

100°

100°

Boundary between eastern and western terranes

Western andesite (Qtz±Calc±Ad±Ill)

Deposits

Ag-Au vein deposit

Basalt-rhyolite

Western andesite (Qtz+Alun±Pyroph±Dick±Kaol)

Ag-Pb-Zn vein deposit Ag-Pb-Zn (Cu) carbonate replacement deposit

Qtz+Alun±Pyroph±Dick±Kaol

30°

Qtz±Calc±Ad±Ill

Other Epithermal Deposits

USA

Mexico

Pacific Ocean

CC

FIG. 17. Map of major epithermal deposits occurring in western United States and Mexico that formed from the middleEocene to the early Pliocene, in comparison to the distribution of coeval rhyolitic deposits (including the San Juan volcanicfield, Colorado, and the Sierra Madre Occidental, Mexico) and the Basin and Range province (Henry and Aranda-Gomez,1992). Most deposits that are portrayed relate to the work of John (2001), Simmons and Albinson (1995), and Albinson et al.(2001), and to some extent the classification schemes used are not comparable (i.e., some deposits with similar characteris-tics are termed differently between regions). In Mexico, Paleozoic-Mesozoic basement rocks are divided into two broadlysimilar packages of rocks comprising eastern and western tectonostratigraphic terranes (Campa and Coney, 1983; Megaw etal., 1988); the western terranes contain rocks of Mesozoic marine and island-arc affinity, whereas the eastern terranes aremade up of a thick succession of Paleozoic to Mesozoic marine carbonates, sandstones, shales, continental red beds and evap-orites that rest on Paleozoic crystalline rocks. The rhyolitic volcanic deposits that cap the Sierra Madre Occidental form avery large continuous rhyolitic ignimbrite sequence (Swanson and McDowell, 1984). The Mexican volcanic belt representsthe modern volcanic arc that trends east-west. Basalt-rhyolite, and western andesite-affiliated epithermal deposits of theGreat Basin are distinguished by John (2001). The Ag-Au and Ag-Pb-Zn epithermal deposits and the Ag-Pb-Zn (Cu) re-placement deposits, based on a classification scheme different from John (2001), are from Albinson et al. (2001) and Megawet al. (1988). Other deposits not discussed in the text are shown for geographical reference. Abbreviations for deposits: AU= Aurora, BD = Bodie, BU = Bullfrog, CC = Cripple Creek; CL = Comstock Lode, C = Creede, F = Fairview, FO = Fres-nillo, GF = Goldfield, GO = Guanajuato, I = Ivanhoe, LC = Lake City, M = Midas, MO = Mogollon, MS = Mulatos, MC =Mule Canyon, ML = McLauglin, N = National, OM = Oatman, PA = Pachuca-Real del Monte, PP = Paradise Peak, RA =Real del Angeles, RM = Round Mountain, SB = San Martín de Bolaños, SL = Sleeper, S = Summitville, TA = Tayoltita, T =Tonopah, TU = Tuscarora, W = Wonder, ZA = Zacatecas.

inclusion studies (Simmons and Albinson, 1995; Albinson etal., 2001, Camprubí et al., 2003) have resolved genetic as-pects that can be correlated further to regional magmatismand tectonism described below.

The epithermal deposits of northern Mexico are divisibleinto two distinct parallel but overlapping metallogenic beltscontaining Ag-Au and Ag-Pb-Zn ores that extend >1,000 kmin a northwest-southeast trend (Fig. 17; Damon et al., 1981;Simmons and Albinson, 1995; Albinson et al., 2001). Some ofthe deposits are middle to late Eocene in age, but mostformed in the Oligocene to Miocene during the Sierra MadreOccidental ignimbrite event when silicic volcanism was bothintense and widespread. The remnant outcrops of this periodare represented by sequences of pyroclastic deposits, manydensely welded with an average thickness of 1 km, which ex-tend over a region 296,000 km2 (Swanson and McDowell,1984). At the same time, epithermal belts formed during ex-tensional tectonism within the southern part of the Basin andRange province, which extends into central Mexico (Henryand Aranda-Gomez, 1992).

The Ag-Au deposits are concentrated in the western belt,commonly occurring in andesitic rocks that comprise thelower part of the Sierra Madre Occidental sequence (e.g.,Tayoltita); their fluid inclusions have salinities of <1 to 7.5 wtpercent NaCl equiv. The Ag-Pb-Zn deposits (e.g., Fresnillo,Zacatecas) are concentrated in the eastern belt and are asso-ciated with the upper part of the Sierra Madre Occidental se-quence comprising felsic pyroclastic deposits. To the east ofthe Ag-Pb-Zn epithermal deposits lies another subparallelmetallogenic belt consisting of Pb-Zn-Ag-(Cu) deposits thatformed at higher temperatures (>300°C), usually close to in-trusive centers and predominantly hosted by carbonate units(Fig. 17; Megaw et al., 1988). Fluid inclusions in the Ag-Pb-Zn deposits generally have higher salinities, ranging from ~5to 23 wt percent NaCl equiv (Simmons and Albinson, 1995;Albinson et al., 2001), with the ore-related salinity at Fresnillobeing ~10 wt percent NaCl equiv compared to the ganguestage of ~2 wt percent NaCl equiv (Simmons et al., 1988).

The Ag-Au and Ag-Pb-Zn epithermal belts extend across amajor crustal suture that separates two sequences of base-ment rocks comprising Paleozoic-Mesozoic tectonostrati-graphic terranes (Fig. 17). Although the presence of evapor-ites in the eastern terrane has been suggested as a possiblesource of high salinities seen in studies of carbonate-hostedPb-Zn deposits (Megaw et al., 1988; Haynes and Kesler,1988), they appear to be missing from the western terrane,which also hosts Ag-Pb-Zn deposits. There is also clear evi-dence from fluid inclusion and stable isotope studies thatmetalliferous brines in these carbonate-hosted Pb-Zn de-posits were derived from magmas (e.g., Sawkins, 1964; Rye,1966), so the contribution to brine formation of evaporitesversus magmas remains unclear. We believe the patterns ofore deposit distribution relative to the regional geologic fea-tures favor a prevailing deep-seated control on mineralizationthat links ore genesis to magmatic processes.

Genetic ConsiderationsEpithermal deposits represent a diverse class of ore de-

posits, and many deposits clearly have different origins, butthe formation of all deposits is dependent on a complex

interplay of physical and chemical factors. Physical factors(e.g., intrusion size and emplacement depth, the permeabilitystructure which evolves during hydrothermal activity, and thelocation of the water table) control heat flow, fluid flux, pres-sure-temperature gradients, flow directions, and sites ofmetal deposition. Chemical factors (e.g., magma composition,crystallization history, influx of meteoric water, and host rockcompositions) affect fluid composition, ligand availability,metal inventory, and metal transport. These factors involvetectonics, magmatism, host rocks, and hydrology, which all acton the overall hydrothermal system to influence the metal ra-tios, mineralogy, and mineral zoning, plus grade, tonnage,shape, and size. Although these factors are known to be im-portant, the timing of processes and the magnitude of theireffects are difficult to measure, so inferences are made byanalogy with active hydrothermal systems.

High metal solubilities and favorable temperature, pres-sure, and host-rock conditions along vertical and horizontalflow paths of several-kilometers distance are required formetals to be delivered to an epithermal environment. More-over, ore textures and discrete occurrences of precious metalssuggest that the critical ore-forming conditions only developepisodically, and ore forms when they are synchronized to op-timize both transport and deposition of metals. In some de-posits we can relate these to pulses of deeply derived metal-bearing solutions rising into the epithermal environment(e.g., Simmons, 1991), whereas in others they appear relatedto shallow fluctuating permeability that governs fluid flow andboiling at the site of mineralization (e.g., Saunders, 1994;Christenson and Hayba, 1995). The limited data regardingthe duration of ore formation indicate that the periods areshort, ranging from a few thousand to a few hundred thou-sand years at most (e.g., Brown, 1986; Hedenquist et al.,1993; Arribas et al., 1995a; Henry et al., 1997; Carman, 2003;Leavitt et al., 2004).

In summary, there are processes within, proximal to, anddistal from epithermal environments that lead to formation ofore deposits, but no one deposit, district, or active hydrother-mal system reveals a unifying framework, because epithermalmineralization is so diverse. We touch on only three out ofseveral aspects relevant to understanding the genesis of ep-ithermal mineralization. These are the chemistry of ore-form-ing solutions, the likely causes of metal deposition, and therole of the water table in controlling the site of ore deposition.

Chemistry of ore-forming solutions

The chemistry of ore-forming solutions, which may also in-fluence gangue and alteration mineral assemblages, evolvesfrom deep to shallow levels (Fig. 1). In both geothermal andmagmatic hydrothermal systems, the proportions of deepfluid (i.e., magmatic versus deeply circulated meteoric wa-ters), followed by the extent of water-rock interaction, are themain factors that influence the chemistry of hydrothermal so-lutions rising into the epithermal environment (Giggenbach,1992a, 1997). Not only do these factors control the pH, theyalso control the oxidation and sulfidation state of hydrothermalsolutions (Giggenbach, 1992a, 1997; Einaudi et al., 2003).

Stable isotope data (Fig. 9) indicate that the origins of theacid solutions responsible for the epithermal gangue assem-blage quartz + alunite ± pyrophyllite ± dickite ± kaolinite can

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be traced directly to shallow intrusions that exsolve a fluid inmuch the same manner as is inferred for the development ofporphyry copper deposits (Hedenquist et al., 1998; Heinrichet al., 2004) and degassing volcanoes (Giggenbach, 1992a,1997). The buoyant magmatic vapor separates from a hyper-saline liquid at depth, rises with little or no water-rock inter-action, and eventually condenses in the epithermal environ-ment. These solutions contain significant quantities of HCl,SO2, and HF, and they become very reactive due to dispro-portionation of SO2, forming H2SO4 and H2S, and dissocia-tion of the acids. Intense acid leaching and formation of vuggyquartz result from reaction with country rock, creating highlypermeable zones that can focus the subsequent throughput ofmetal-bearing liquids (White, 1991; Arribas, 1995; Heden-quist et al., 1998; Heinrich, 2005). The early acidic solutionsare progressively neutralized through hydrolytic alteration ofalumino-silicate minerals (e.g., feldspars) accounting for thesharp outward zoning in alteration assemblages (Fig. 7). Inmany cases, however, a subsequent metal-bearing liquid doesnot reach the level of advanced argillic alteration, leaving itbarren of ore mineralization. Such barren zones are associ-ated with porphyry copper deposits creating topographicallyprominent features known as lithocaps (Sillitoe, 1995b). Min-imal water-rock interaction in ascending metal-bearing solu-tions is also reflected by the high-sulfidation state of Cu sul-fide minerals. The relatively oxidized nature of thesesolutions and their evolution to a high-sulfidation state oncooling from magmatic to epithermal temperatures resultsfrom buffering with aqueous H2S and SO2 (Giggenbach,1992a, 1997; Rye, 1993, 2005; Einaudi et al., 2003). Increas-ing duration of water-rock interaction leads to intermediate-sulfidation state assemblages, which commonly form late inthe parageneses of enargite-bearing deposits (Einaudi et al.,2003). Such changes in fluid chemistry appear essential to op-timizing gold transport into the epithermal environment(Heinrich, 2005). Note that the composition of fluids causingquartz + alunite ± pyrophyllite ± dickite ± kaolinite alterationdiffers from the composition of fluids transporting the metals,even if both fluids have a common heritage, relating to thesame intrusive stock (White, 1991; Arribas, 1995; Hedenquistet al., 1998; Heinrich, 2005).

Stable isotope data (Fig. 9) indicate that the origins of near-neutral pH solutions responsible for the epithermal gangue as-semblage quartz ± calcite ± adularia ± illite originate from deepcirculation of meteoric water. Based on the geothermal systemanalogy, we infer that the fluid in such systems is influenced byvariable input of magmatic components (e.g., nil to ~10% mag-matic H2O) followed by extensive water-rock interaction (Fig.1). The resulting solution is near-neutral pH and largely in equi-librium with a propylitic alteration assemblage. Minerals such asK-mica and K-feldspar buffer the pH, and iron-bearing phases(silicates and sulfides) buffer the oxidation state (Giggenbach,1992a, 1997), with the formation of low- to intermediate-sulfi-dation–state minerals possibly reflecting the nature of the par-ent magma (John, 2001; Einaudi et al., 2003; Sillitoe andHedenquist, 2003). Waters with high δ18O values suggest thatformation waters may have been present in some deposits (e.g.,McLaughlin, United States; Sherlock et al., 1995).

The salinities of ore-forming solutions (Fig. 8) vary widely.In individual deposits, a salinity range of a factor of 2 or 3 can

be attributed to boiling and/or mixing (e.g., Hedenquist andHenley, 1985b; Hayba, 1997), although very locally a range often times can result due to boiling and evaporation in an iso-lated fracture (Simmons and Browne, 1997). However, thesesame processes cannot account for the differences in salini-ties, for example, between Ag-Pb-Zn (5–23 wt % NaCl equiv)and Au-Ag (<1–7.5 wt % NaCl equiv) in northern Mexico(Simmons and Albinson, 1995; Albinson et al., 2001). More-over, the more saline hydrothermal solutions (i.e., >5 wt %NaCl equiv) have no modern counterpart in analogous activehydrothermal systems, and the stable isotope data indicatethat salinities reflect fluid pulses having significant magmaticinput (Simmons, 1991, 1995; Albinson et al., 2001). The oc-currence of evaporites within the host-rock stratigraphy mayin some cases account for high fluid salinities, as seen in theSalton Sea geothermal system (e.g., McKibben and Hardie,1997) and in some Mexican Ag-Pb-Zn deposits, discussedabove, but these appear to be the exception rather than therule. Thus, the range of salinities observed in fluids that formdeposits dominated by intermediate- versus low-sulfida-tion–state sulfides hosted by quartz ± calcite ± adularia ± il-lite appears to reflect different sources. For fluids that formhigh- to intermediate-sulfidation–state sulfides hosted byquartz + alunite ± pyrophyllite ± dickite ± kaolinite, moder-ate- to high-salinity fluid inclusions reflect the history of fluidexsolution during crystallization of underlying magmas (e.g.,Hedenquist et al., 1998; Heinrich et al., 2004; Heinrich,2005), although most salinities are low, <5 wt percent NaClequiv.

In summary, the difference between acidic magmatic hy-drothermal solutions and near-neutral pH chloride waters isrelated to (1) the nature and depth of the underlying intru-sion, (2) the resulting length of the fluid-flow path separatingthe epithermal environment from the deeper parts of the sys-tem, (3) the degree to which fluids are able to ascend freelyvia open vertical conduits, and (4) related to the last point, theextent of water-rock interactions as fluids ascend to the sur-face. These factors might explain how in some deposits (e.g.,Chinkuashih-Chiufen, Taiwan; Lepanto-Victoria, Philippines)ores produced in the two contrasting alteration environments(Fig. 1) can form side by side within the same broad period ofhydrothermal activity, as also seen in some active hydrother-mal systems (e.g., Hedenquist and Aoki, 1991; Reyes et al.,1993). What remains puzzling is why some districts or beltsare dominated by ore deposits associated with quartz ± calcite± adularia ± illite gangue (e.g., Hauraki goldfields), some aredominated by ore deposits associated with quartz + alunite ±pyrophyllite ± dickite ± kaolinite gangue (e.g., Yanacocha,Peru), and others have both types of ore deposits (e.g., theGreat Basin in the western United States).

Metal transport and deposition

Metal transport and deposition and the formation of ore-bodies over a restricted vertical interval, a few hundred me-ters maximum, are among the most important processes af-fecting ore genesis in epithermal deposits. Gold and silvertransport in hydrothermal solutions is caused mainly by bisul-fide complexes, whereas the base metals and a component ofthe silver are transported by chloride complexes (e.g., Sewardand Barnes, 1997). The metal contents of deep fluids from

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some geothermal systems have been analyzed from directsampling (Brown and Simmons, 2003) and deduced frommetal precipitation in surface pipes of geothermal wells(Brown, 1986; McKibben and Hardie, 1997). However, dataregarding the metal contents of epithermal solutions is lim-ited and much more is known about the causes of metal de-position in epithermal deposits, especially from near-neutralpH solutions, based on abundant mineralogical evidence re-flecting processes associated with ore formation. This is im-portant because regardless of how much metal might havepassed through a hydrothermal system, ores will not form un-less there is a mechanism for efficient metal deposition.

The most influential agents of metal precipitation are fo-cused fluid flow along with some combination of boiling andmixing, as indicated by the occurrences of ores in zones ofhigh paleopermeability, deductions from fluid-mineral equi-libria, fluid inclusion and mineralogical evidence, and obser-vations in epithermal deposits and active geothermal systems(e.g., Buchanan, 1981; Brown, 1986, 1989; Heald et al., 1987;Seward, 1989, Christenson and Wood, 1993; Saunders, 1994;Cooke and Simmons, 2000; Cooke and McPhail, 2001;Berger et al., 2003). Phase separation due to boiling causes asharp loss of H2, H2S, and CO2 that lowers precious and basemetal solubility where sulfide and chloride complexes domi-nate, and it also causes an increase in the pH and oxidationstate, largely due to the loss of CO2 and H2, respectively. Boil-ing is a highly efficient mechanism which removes most goldand silver from solution, as clearly illustrated by gold and sil-ver deposits in surface pipes mentioned earlier (Brown,1986). Boiling also causes precipitation of adularia, platy cal-cite, and colloform-banded, amorphous silica (Simmons andBrowne, 2000b), and these are found in many epithermal veinores noted above. Many barren veins also show these sametextures, suggesting that although conditions conducive tometal precipitation existed (i.e., boiling), metals were deficientin the hydrothermal solution. There are no known modern ana-logues for precious metal precipitation in a magmatic hy-drothermal system, but chemical arguments and fluid inclusiondata support the importance of boiling (Cooke and Simmons,2000; Deyell et al., 2004), as does the occurrence of ores in hy-drothermal breccias of explosive origin. The development ofsteam-heated waters and corresponding hydrolytic alteration inshallow and near-surface environments corroborates that boil-ing is common in epithermal environments hosted by bothmagmatic hydrothermal and geothermal systems.

Mixing between fluids of different compositions is anotherviable mechanism of precious metal precipitation, as sup-ported by some fluid inclusion data (e.g., Robinson and Nor-man, 1984; Mancano and Campbell, 1995; Hayba, 1997), ex-tensive stable isotope data (Fig. 9), and numerical simulations(Reed and Spycher, 1985; Plumlee, 1994). There are alsocases where mixing has not produced metal deposition (e.g.,Simmons and Browne, 2000b); however, in cases where metaldeposition occurred, it probably resulted mainly from dilu-tion and cooling and, to a lesser extent, from changes in oxi-dation state and pH.

For either mixing or boiling to occur and cause ore deposi-tion, a favorable hydrological setting must exist. Boiling re-quires sharp temperature-pressure gradients and a free fluidpath to the surface (e.g., Simmons and Browne, 2000a, b),

whereas mixing requires sustained interaction between fluidsof different compositions and/or temperatures (e.g., Hayba,1997), preferably in a turbulent environment. But not all ep-ithermal deposits can be easily explained as simple productsof boiling and/or mixing. Among these are the large gold de-posits that contain telluride-bearing ores over long vertical in-tervals (>500 m) in association with alkaline igneous rocks(Emperor, Fiji; Porgera, Papua New Guinea; Cripple Creek,United States). For these, mechanisms of metal depositionseem complex, involving a combination of factors includingboiling, mixing, and water-rock interaction (Ronacher et al.,2004), and possibly reactions involving condensation of Te-bearing magmatic gas into metal-bearing solutions (Cookeand McPhail, 2001).

The role of the water table

The regional water table controls the hydrostatic pressuregradient in subaerial hydrothermal systems (Henley, 1985).Thus, the elevation of the water table and its shift with timerelative to the land surface play an important role in dictatingthe vertical position of the epithermal environment (e.g.,Steven and Eaton, 1975; Henley and Ellis, 1983; Henley,1985; Sillitoe, 1993b). The position of the water table is influ-enced by topography and climate, so that in steep or arid ter-rains there may be as much as several hundred meters be-tween the water table and the land surface, in contrast to flatand wet terrains, where the two closely coincide. Steep ter-rains also influence the regional hydraulic gradient, which in-duces lateral flow and mixing in the epithermal environment,potentially conducive to metal deposition (e.g., Henley andEllis, 1983; Hayba, 1997). Despite the dynamic nature of sur-face changes in volcanic terrains, the effects of the water tableon mineralization are not commonly documented. There arerelatively few examples where the water table and its evolu-tion have been interpreted and substantiated as integral toore genesis (Simmons, 1991; Sillitoe, 1994; Ebert and Rye,1997; Bissig et al., 2002; John et al., 2003; Wallace, 2003).

The water table rises and falls in volcanic arcs under the in-fluence of uplift, subsidence, erosion, volcanic eruption, andlake drainage. These effects range from local (<1–10 km2) toregional (>10,000 km2) in extent. Whereas the minimumamount of time required to form an epithermal orebodymight be a thousand years or more (Henley, 1985; Brown,1986; Hedenquist et al., 1993), large shifts in the water table(>100 m) can occur on the scale of hours to months, for ex-ample, due to volcanic eruption (cone building or caldera for-mation), sector collapse, or breakout flooding causing cata-strophic drainage of a lake-filled depression (e.g., Goff et al.,1989; Simmons et al., 1993; López and Williams, 1993;Manville et al., 1999). In such examples where the water tablefalls rapidly, the accompanying pressure drop may trigger hy-drothermal eruption, brecciation, and precious metal deposi-tion, whereas in examples of progressive erosion, a fall in thewater table may telescope alteration styles and orebodies(e.g., Simmons, 1991; Sillitoe, 1994, 1999). By comparison,water table changes induced by steady regional uplift of a fewmm/year require several thousand years or more to have acomparable effect, although this can be well within the lifes-pan (<100,000 yr) of a single hydrothermal system (e.g.,Reyes, 1990; Bignall and Browne, 1994).

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The overall effect of shifting the water table up or downduring hydrothermal activity is to extend or contract the dis-tance separating rocks altered and mineralized in deep andshallow environments and, if precious metal mineralization isforming at the same time, to change the vertical distributionof ores. Evidence of the position of the water table is deducedfrom fluid inclusion data and shallow alteration patterns(blankets of steam-heated advanced argillic alteration, silicasinter). In regions where the water table is rising (e.g., due toregional subsidence or damming of a drainage), hotter alter-ation assemblages prograde on to cooler alteration assem-blages and silica sinters may become stacked in the strati-graphic record (e.g., Hasbrouck Mountain, United States,Graney, 1987; Drummond basin, Australia, Cuneen and Silli-toe, 1989). In regions where the water table drops, cooler al-teration assemblages retrograde on to hotter assemblages andmineralization is telescoped. This latter situation may explainoccurrences of epithermal- over porphyry-style mineraliza-tion (e.g., Sillitoe, 1999) or, in the very rare case, in the vicin-ity of batholith intrusions (Kesler et al., 2004).

Implications for ExplorationEpithermal deposits are variable in size, shape, and grade,

and commonly the ore zones are not exposed. These charac-teristics make them elusive to find and a challenge to explore(e.g., Sillitoe, 1995, 2000). Although some deposits are largeand continuous, many are not (Fig. 3). Efficient explorationthus requires integration of all available geological, geochem-ical, and geophysical data with a good understanding of de-posit characteristics and ore-forming processes, plus a will-ingness to drill targets generated from these data. Epithermaldeposits have the benefit that there are many features, as dis-cussed above, that provide valuable information on erosionlevel and mineral zonation.

There have been many exploration successes since the re-vival of interest in epithermal deposits in the late 1970s. Somehave resulted from reevaluation of the potential for low-gradebulk mining of deposits originally exploited as high-grade de-posits (e.g., Martha Hill, New Zealand; Round Mountain,United States), and others have resulted from discoveries ofnew deposits in known mining districts (e.g., McLaughlin andSleeper, United States; Golden Cross, New Zealand;Hishikari, Japan; San Cristobal, Bolivia). Some have been theresult of expanding exploration into previously unexploreddistricts (e.g., Ladolam, Papua New Guinea; Mount Muroand Kelian, Indonesia; El Indio, Chile; Cerro Vanguardia, Ar-gentina). The discoveries of epithermal deposits in areas orig-inally being explored for other types or styles of mineraliza-tion (e.g., Midas, United States; El Peñon, Chile; Victoria,Philippines; Nevada, Chile; Veladero, Argentina; Sillitoe,1995, 2000) highlight the critical importance of explorersbeing familiar with the key characteristics of the differentstyles of epithermal mineralization so they can recognize itssignificant traces, even if exposed in only small areas.

The first objective in exploration for epithermal deposits isto choose the favorable regions to explore and then to narrowthis extensive prospective region to a manageable area for de-tailed exploration, and finally to define targets for drill testing.Once a potentially economic intersection has been achieved,further work is mostly resource definition and evaluation,

rather than true exploration, although it is commonly con-ducted by the same staff and many of the same skills are re-quired. The distinction here is between exploring for a de-posit and assessment of a deposit.

There are few technical papers that address exploration is-sues directly (Sillitoe, 1995a, b, 2000). White and Hedenquist(1990, 1995) and Hedenquist et al. (2000) described and dis-cussed attributes of epithermal gold deposits that are usefulfor exploration. More specific geochemical and geophysicalaspects were addressed by Allis (1990), Clarke and Govett(1990), Irvine and Smith (1990), Ellis and Robbins (2000),and Wright and Lide (2000). Table 6 shows the main goalsand methods used at different scales of the explorationprocess. Some exploration programs cover all scales, whereasothers begin with a relatively well-defined target. In all cases,the aim is to reduce the area, and then the volume of rockbeing explored, with the goal of ultimately defining an eco-nomic deposit.

Exploration techniques can be conveniently classified asgeological, geochemical, and geophysical in emphasis. In gen-eral, the geologic techniques are those that geologists dothemselves and are dependent on recognition and mapping ofgeologic features. Geochemical techniques are those that in-volve determining the concentration of various chemical ele-ments. Geophysical techniques involve measurement of geo-physical parameters that are subsequently inverted toproduce a geologic interpretation.

Geologic techniques

The aim is to use geologic characteristics to identifyprospective areas for more detailed exploration. Such areasare indicated by the occurrence of known deposits andprospects, either in the area chosen for exploration or in otherareas that are thought to be geologically similar. Other favor-able features include hydrothermal alteration of appropriatestyles and dimensions (Figs. 4, 6, 10) and/or encouraging geo-chemical signatures. Apart from prospecting and groundmapping, remotely sensed images and spectral data can beused to recognize areas of hydrothermal alteration in suitableregions (especially arid, poorly vegetated areas). Traditionalinterpretation of aerial photographs is an effective techniquethat has largely fallen into disuse. Mapping of structures andtheir analysis nonetheless remain a key guide to ore.

Even at a regional scale, prospecting methods can play acrucial role when exploring for epithermal deposits. In mostdeposits the main ore zone consists of siliceous rocks (veins orsiliceous alteration) that are very resistant to erosion. Miner-alized clasts are likely to be transported into streams and canbe recognized by reconnaissance surveys at an early stage,confirming the potential of an area selected on conceptualgrounds. Barren siliceous material that forms above and distalfrom epithermal mineralization (e.g., silica sinter and opalineblankets; Fig. 5) can also point to paleohydrothermal activity.

As exploration becomes more detailed, the emphasis shiftsfrom looking for broad areas of hydrothermal alteration, tofinding zones of focused fluid flow and zones of ore deposi-tion. Mapping of the host geology, structures, vein texturesand alteration mineralogy are all important. The availability ofaffordable field infrared spectrometers has changed the ap-proach to exploration of epithermal deposits, because they

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allow inexpensive and rapid mapping of fine-grained hydro-thermal alteration. These studies are applicable at all scales ofexploration.

Geochemical techniques

In gold exploration by far the most important element ana-lyzed is gold, although other elements are commonly associ-ated with gold deposits and they can also be used in explo-ration. In epithermal gold deposits the suite of associatedelements is large and includes Ag, Cu, Zn, Pb, As, Sb, Bi, Se,Te, Tl, Mo, W, Sn, and Hg.

Traditional panning for gold, as practiced by prospectorsfor hundreds of years, has mixed success for epithermal tar-gets. The discovery of the Kelian deposit, which resulted fromexploration upstream of known alluvial gold deposits, is a no-table example (van Leeuwen et al., 1990). In deposits associ-ated with quartz ± calcite ± adularia ± illite, the gold is typi-cally in the form of electrum, which is more susceptible todissolution in ground water than native gold, especially in re-gions with acidic or saline ground waters (e.g., Webster andMann, 1984; Webster, 1986). In deposits associated withquartz + alunite ± pyrophyllite ± dickite ± kaolinite, the goldis partly native, but it is typically fine grained and much of thegold is refractory, contained in sulfide minerals. Conse-quently, epithermal deposits commonly produce less alluvialgold than might be expected from their gold content.

Regional-scale geochemical surveys typically use stream-sediment geochemistry, commonly conventional (–80 mesh),or low-density bulk cyanide leach (BLEG); pan-concentrategeochemical surveys also have been employed. As follow-up,or for smaller areas, conventional surveys are more common.Once areas of interest have been defined by early surveys, fol-low-up is commonly by grid-based soil surveys. In rugged ter-rain where grids would be impractical, ridge and spur sam-pling is commonly used.

When geochemically anomalous areas have been defined,but the source is not exposed because of cover by soil orscree, trenching is commonly done to expose bedrock. Thisallows detailed geologic mapping as well as rock-chip sam-pling to define the source of the anomaly. Once the sources ofanomalies have been located, and the geology has beenmapped, exploration requires testing of targets at depth bydrilling.

Geophysical techniques

Advances in geophysical techniques over the last 10 yr haveimproved our ability to explore for epithermal deposits. In ex-ploration of geothermal areas, both resistivity and gravimetricsurveys have been employed (Allis, 1990; Bibby et al., 1995).In geothermal areas, the low resistivity response is related tothe hot water in the rocks and/or clay alteration, and thus it isnot as effective as in inactive systems which host the epither-mal deposits. Gravimetric surveys detect the increase in den-sity that results from hydrothermal minerals deposited in per-meable volcanic units and can also be used to locatetopographic highs on the underlying basement rocks.

Intense hydrothermal activity commonly destroys mag-netite in the host rocks by converting it to pyrite. Large-scalemagnetic surveys can detect these areas of demagnetization(Irvine and Smith, 1990); however, the lack of discernible fea-tures inside these areas means that a much larger area needsto be surveyed so that structures can be detected outside thealtered area and traced into the area of interest (e.g., Hilde-brand et al., 2001).

At a more detailed scale, several geophysical techniquescan be useful in defining the position of zones of interest atdepth below surface anomalies. The extensive zones of dis-seminated pyrite that characterize most deposits hosted byadvanced argillic alteration are suitable for induced polariza-tion (IP) surveys. However, some deposits associated with

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TABLE 6. The Main Objectives and Methods Used at Different Scales of Exploration for Epithermal Deposits

Scale and objective Look for Methods used (importance)

Regional (×00–×000 km2) Known deposits and/or prospects Literature review, prospecting (high); Locate favorable belt; Alteration zones Literature review, prospecting, satellite and aerial define prospective areas photograph studies (high);

Geochemical anomalies Literature review, prospecting, regional-scale surveys (high);Geophysical anomalies Literature review, reevaluation of regional magnetic,

EM and gravity surveys (moderate)

District (×00–×00 km2) Known deposits and/or prospects Literature review, prospecting, aerial photo studies (high)Define prospects Alteration zones Prospecting, ground surveys, satellite and aerial photograph

studies (high)Geochemical anomalies Stream and soil geochemical surveys (high)Geophysical anomalies Large-scale magnetic surveys (low)

Prospect (<1–100 km2) Veins, mineralized structures and/or zones Mapping (incl. trenching) (high)Alteration zones Mapping with IR spectrometry (incl. trenching) (high)Geochemical anomalies Detailed soil and rock chip surveys (high);Geophysical anomalies IP or resistivity surveys (moderate)

Deposit (mostly <1 km2) Veins, mineralized structures and/or zones Mapping, drill logging (very high)Alteration zones Mapping, drill logging with IR spectrometry, XRD (very high)Geochemical anomalies Detailed geochemistry of drill samples or underground

workings (very high)Geophysical anomalies Resistivity (moderate)

alteration that formed from near-neutral pH solutions alsohave abundant disseminated pyrite, and all deposits can un-derlie a blanket of pyritic clay-carbonate alteration, whichformed near the surface in the deeper parts of steam-heatedzones.

High-resistivity siliceous zones are intimately associatedwith mineralization in the epithermal environment, both invuggy and massive quartz as well as in quartz veins. Suchzones are amenable to detection by resistivity surveys, such ascontrolled-source audio-frequency magneto-telluric (CSAMT;Irvine and Smith, 1990), as well as by IP. The limitation onsuch surveys is that the high-resistivity target zones are com-monly narrow in vein-related deposits, and associated withother zones that have low resistivity (e.g., clay zones) and highchargeability (due to sulfides), making their interpretationvery difficult. Despite this problem, geophysics is the only ex-ploration technique that allows measurement of subsurfacegeologic properties prior to drilling, and its role in explorationfor epithermal deposits is likely to increase, particularly as thetechniques are refined.

Concluding RemarksEpithermal deposits formed in the shallow part of once-ac-

tive hydrothermal systems. They are a diverse class of ore de-posits, containing ores of differing metal compositions, min-eralogies, and origins. We take the perspective thatepithermal deposits are best examined in terms of their com-mon gangue mineral assemblages. This is based on empiricalobservations, but it is underpinned by our understanding ofthe conditions under which these gangue minerals form interms of temperature, pressure, fluid composition, and iso-topic composition.

In the past few years, the importance of magmas and tec-tonic setting, and their influence on metal inventories, sulfida-tion state of ore minerals, and styles of alteration and mineral-ization have been discussed (John, 2001; Sillitoe andHedenquist, 2003). In some places, such as the Great Basin ofthe western United States (John, 2001), established relation-ships have helped to develop a genetic framework for epither-mal ore formation. However, there are other places wherethese relationships are difficult to establish, and it remains tobe seen how often correlations can be made. Continuity in therecord of igneous activity and the ability to identify time mark-ers that match ore formation to intrusive or volcanic eventswill help to strengthen the link between magmatism and stylesof epithermal mineralization. In the case of the high- to inter-mediate-sulfidation–state ores associated with quartz + alunite± pyrophyllite ± dickite ± kaolinite gangue, precious and basemetal mineralization is intimately associated with the crystal-lization of igneous intrusions and exsolution of magmatic flu-ids, and knowledge of their effects will improve the under-standing of genetic processes beneficial to exploration.

The same strong link is not as obvious in many intermedi-ate- and low-sulfidation ores hosted by quartz ± calcite ±adularia ± illite gangue, because the distance separating ig-neous intrusions and orebodies is longer (typically severalkilometers), and the fluids rising from the base of the con-vection cell are subject to greater degrees of water-rock in-teraction that can mask magmatic effects. There are also theeffects of the local epithermal environment, for example,

boiling and mixing, which may cause metal deposition,changes in the sulfidation state, and influence the sequenceof mineral precipitation. Consequently, without knowing thecompositions of deep fluids in detail (specifically CO2, H2S,SO2, H2, Cl, and metal contents), the distinction between oreswhose mineralogy is controlled by igneous intrusion and oreswhose mineralogy is controlled by the processes in the ep-ithermal environment is blurred.

Important avenues of future research of epithermal de-posits include: (1) continued application of high resolution40Ar/39Ar dating to resolve the timing of ore formation relativeto the age of host rocks, and the history of igneous activity,and to understand the duration and frequency of mineralizingevents; (2) quantitative determination of the compositionsand metal contents of ore-bearing fluids from microanalysisof fluid inclusions as has been used to study porphyry deposits(e.g., Ulrich et al., 2001); this should lead to a great improve-ment in understanding epithermal deposits because of thepotential to characterize metal contents of fluids related tospecific events; (3) sampling and trace metal analysis of fluidsin active systems (e.g., Brown and Simmons, 2003) to resolvethe diversity of fluid types, especially of deep origin, and de-termine their capacity to transport metals in different geo-logic settings; this in conjunction with results from inclusionfluid analyses will contribute to quantitative models of metaltransport and deposition that may increase the understandingof how trace element dispersion halos relate to ore-formingprocesses; and (4) new exploration techniques (including geo-chemistry and geophysics) that improve discovery of orebod-ies concealed beneath cover rocks; the tops of many large ep-ithermal deposits remain to be recognized, and they will befound both in virgin territory and in known districts whoselimits have yet to be defined.

AcknowledgmentsA number of people generously provided data on epither-

mal deposits, including Regina Baumgartner, Bob Brath-waite, Tony Christie, Jillian Exton, Mark Fisher, Tony Longo,Jeff Mauk, Sachihiro Taguchi, and Steve Turner. We thankShane Ebert, Patrick Browne, and Bob Brathwaite for theirreviews on an early draft of this manuscript. We especiallythank Jeff Hedenquist and John Thompson for their tirelessefforts and their technical and editorial comments on the sub-sequent drafts, which greatly improved this manuscript. MarkSimpson deserves special mention for his help in compilingore deposit data and drafting of figures. This work was sup-ported by a grant from the Foundation for Science, Researchand Technology to SFS.

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