1995 Arribas

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A .

A rribas, Jr.

Table

Nin

Fig. 1

1

2

3

4

5

6

7

8

9

10

11

12

13

14

15

16

17

18

19

20

21

22

23

24

25

26

27

28

29

30

31

32

33

34

35

36

37

38

39

40

41

42

43

44

45

46

47

48

49

50

1. Principal high-sulfldation deposits or

Deposit

documented prospects ordered geographically

References

Asia Australasia

Dobroyde, Australia

Rhyolite Creek, Australia

Temora, Australia

Peak Hill, Australia

ML Kasi, Fiji

Wafi

River,

PapuaNewGuinea

Nena, Papua

New

Guinea

Motomboto, Indonesia

Nalesbitan, Philippines

Lepanto, P hilippines

Chinkuashih, Taiwan

Zijinshan, C hina

Seongsan & Ogmaesan, South Korea

Nansatsu (Iwato, Akeshi Kasuga), Japan

Yoji, Japan

Teine,

Japan

Akaiwa, Japan

Mitsumori-Nukeishi, Japan

W h i t e ral.(1995)

Raetz Partington (1988)

Thompson

e tal.

(1986)

Cordery (1986), Harbon (1988), M asterman (1994)

Turner (1986)

Leach Erceg (1990), Ercegetal. (1991)

Asami & Britten (1980), Halletal.(1990)

Perelld (1994)

Sillitoe

e tal.

(1990)

Gonzalez (1959), Garcia (1991), Arribasetal.(1995b)

Huang (1955), Hwang Meyer (1982), Tanetal. (1993)

Zhangetal. (1994)

Yoon (1994)

Izawa Cunningham (1989), Hedenquistet

al.

(1994a)

Yui&Matsueda(1994)

Ito (1969)

Akamatsu & Y ui (1992), Akamatsu (1993)

Aoki Watanabe (1995)

North

Central America

Northwestern VancouverIsland,Canada

Goldfield, Nevada

Paradise

Peak,Nevada

Summitville, Colorado

Red Mtn-Lake City, Colorado

Red Mtn-Silverton, Colorado

Mulatos, Mexico

Pueblo Viejo, Dominican Republic

Panteleyev Koyanagi (1994)

Ransome (19 07,190 9), A shley (1974), Vikre (1989)

Johne tal. (1991), Sillitoe

Lorson (1994)

Steven Ratte" (1960), Stoffrcgen (1987), Rye (1993)

Bove

etal.

(1990), Rye (1993)

Burbank (1941), Fisher

and Leedy

(1973)

Staude(1994)

Munteane tal.(1990), Russell Kesler (1991)

South America

Julcani, Peru

Castrovirreyna, Peru

Ccarhuarso, Peru

San Juan de L ucanas, Peru

Cerro

de Pasco, Peru

Colquijirca, Peru

Sucuitambo, Peru

Laurani, B olivia

Choquelimpie, Chile

Guanaco, Chile

El H ueso, Chile

Esperanza, Chile

La Coipa, Chile

Nevada Sancarron, Chile

El Indio-Tambo, Chile

La Mejicana-NevadosdelFamatina, Argentina

Petersenet

al.

(1977), Deen (1990), Rye (1993)

Vidal Cedillo (1988)

Vidal ef a/. (1989)


Graton

Bowditch (1936), E inaudi (1977)

Vidal

etal.

(1984)


Murillo etal. (1993)

GiOpper etal.

(1991)

Puigetal.(19 88), Cuitifioetal.(1988)

Sillitoe (1991a)

Vila (1991), Moscosoe tal.(1993), Cuitifioetal. (1994)

Oviedoetal.(1991), Cecioni

Dick (1992)

Siddeley Araneda (1990)

Siddeley Araneda (1986), Jannasetal.(1990)

Losada-Calderon McPhail (1994)

Europe

Rodalquilar, Spain

Furtei-Serrenti, Sardinia

Spahievo, Bulgaria

Chelopech, Bulgaria

Western Srednogorie region, Bulgaria

Bor, Yugoslavia

Lahoca, Hungary

Enasen, Sweden

SSnger-von Oepenetal.(1989), Arribasetal.(1995a)

Ruggieri (1993a,b)

Velinovrtc/.(1990)

Bogdanov (1982,1986)

Bogdanov (1982), Velinov Kanazirski (1990)

Jankovic etal.(1980), Jankovic (1982)

Baksa (1975,1 986), First (1993)

HaUberg(1994)

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High-sulfidation Epithermal Deposits

m

WT

C ^ I U - 1 0 Western

i f l P J 5V -9 V

*

lc

jk /S' fy g / Pacific

- - T V ?

Figure 1. Worldwide distribution of high-

8/11/2019 1995 Arribas

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A .

A rribas, Jr.

Table

2.

Main geological characteristics of

14

selected high-sulfidation epithermal deposits

Deposit/district,

location

Motomboto,

Indonesia

Nalesbitan,

Philippines

Lepanto,

Philippines

Chinkuashih,

Taiwan

Zijinshan,

China

Nansatsu,

Japan

Summitville ,

Colorado

Goldfield,

Nevada

Paradise Peak,

Nevada

Pueblo Viejo,

Dominican Rep.

Julcani,

Peru

El Indio,

Chile

La Mejicana & N e-

vados del Famatina,

Argentina

Rodalquilar,

Spain

Age

(Ma)

1.9

Pliocene

1.5-1.2

1.3-1.0

-94

5-3.5

22.5

21

19-18

- 1 3 0

9.8

13-8

4.0-3.6

11-10

Metals,

(tonnes)

1

Cu, Au, Ag

60 ,000

t

Cu .

4 i

A u , 1 8 0 t A g ( c )

Au

15t Au (c)

Cu, Au,Ag

900 ,000 iCu,

1201 Au(c )

Au, Cu, Ag

92 t Au, 183 t Ag,

120 ,000 tCu(p)

Cu , Au

> 1 0

t

Au (c)

Au

18

t

Au (p) + 18

t

Au reserves

Au, Cu,

Ag

1 7 l A u

Au (Ag, Cu)

13 0tAu, t 43 Ag,

37,000 Cu(p)

Au, Ag, Hg

47 t Au, 1255 Ag

45 7lHg (p)

Au ,

Ag

> 6 0 0

t

Au

(p;

Sillitoe.

1993)

Ag, Cu, Pb, Au,

W, Bi, Zn

Au, Ag, Cu

- 1 4 0 tAu,

- 1 , 1 0 0 t A g ( c )

Cu ,

Au ,

Ag

> 1 0 - 1 5 t A u ( c )

Au

10 l Au (p)

Local volcanic

setting

Central-vent

volcano

Small central-

vent volcano

Diatreme

complex

Dome complex

Dome along

caldera margin?

Small volcanos

inacaldera?

Dome along

preexisting

caldera margin

Domes along

preexisting ring

fracture

Within or close

to a central-vent

volcano

Maar-diatreme

complex

Dome complex

aroundacentral

diatreme

Stratovo cano(?)

inearliercaldera

Dome complex(?)

Caldera margin

Principal host

rocks

Dae d ome, ands/dac/rhy

flows, pyr and volx

Ands pyr

+

flows

Ands/dac vol,

Miocene+older

volx+ metavol

Dae vole

Miocene sed

Jurassic granite,

Cretaceous dac

porpyhry+pyr

Ands pyr, flows

+

volx

Qu-lalite porphyry

Miocene andesite

Composite welded

tuff,

volx

+

ands flows

Maar sed

+

basaltic

vol (spilite)

Dactorhyodacit ic

domes and tuffs

Dac , rhy pyr;

dac

+

ands vol

Paleozoic seds+

granites. Pliocene

intrusive dacite

Ands to rhy pyr flows,

collapse bxs+domes

Genetically

related rocks

Dioritic, qtz-

dioritic stocks

None observed

Qtz-diorite

porphyry

Dacite domes

and flows

Not reported

HorWende ands

(Middle Voles)

Qtz-monzonite

porphyry

Andesite

And/dacvd

CAbimodal

(Rhy

+

basalt)

volcanic suite

Dac/rhyodacitic

porphyry

CA vol

Dac/rhyodaoic

porphyry

stocks

Ands flows

+ dykes

Time

between host

rock & deposit

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Table 2 (continued)


Deposit/district,

location

Motomboto,

Indonesia

Nalesbitan,

Philippines

Lepanio,

Philippines

Chinkuashih,

Taiwan

Zijinshan,

China

Nansatsu,

Japan

Summitville ,

Colorado

Goldfield,

Nevada

Paradise Peak,

Nevada

Pueblo Viejo,

Dominican Rep.

Julcani,

Pem

El Indio,

Chile

La Mejicana & N e-

vados del Famatina,

Argentina

Rodalquilar,

Spain

Control on mineralization

Contact between dom e and

volcanic

rock,

steep fault

Steep strike-slip fault

Major steep + minor faults,

diatreme contact, unconfor

mity, permeable layers

Steep normal faults +

their intersections,

bedding planes

Steep strike-slip fault

zones + contact of

volcanic vent

Steep fractures + permeable

pyroclastic layers

Steep radial fractures +

dome contact

Moderately + shallow

dipping faults & fissures

Steep faults and permeable

pyroclastic layers

Diatreme ring fault +

permeable layers

Steep fractures

Steep normal faults

Local faults

Caldera ring faults +

normal local faults

Vertical ext

ent of epith

ore(m)2

250

ISO

500

800

600(7)

< 1 5 0

250

400

3 0 0

< 1 5 0

Relation to

porphyry sy stem

Porphyry Cu-Au

prospects nearby, age

within 1.0 m.y.

Proposed,

none known

Above + adjacent

same age porphyry

Cu-Au deposit

None known

None known

None known

Intrusion-centered

sericitic, low grade

stk mineralization

None known

Sericitic, stk Au

mineralization (East

Zone)

None known

None known

Porphyry Cu-Mo

mineralization

nearby

HS ore at Nevado del

Famatina is a part of a

porphyry Cu prospect

None known

References

Pere l l6 ( l94)

Si l l i toer ta / . ( 1990)

Garcia (1991),

Arribasrta/. (1995b)

Huang (1955),

Ta n

etal.

(1993)

Ren era/. (1992),

Zhangetal. (1994)

Izawa & Cunningham (1989),

Hedenquist

etal.

(1994a)

Steven & Ratti (1960), Menhert

etal

(1973 ), Stoffregen (1987 ),

Rye (1993) Gray & Coolbaugh

(1994)

Ransome (1909). Ashley (1974),

Ashley & Silberman (1976),

Vikre (1989, written commun.

1995)

John

et al.

(1991),

Sillitoe & Lorson (1994)

Russell &Kesler (1991),

Mumeanero/. (1990)

Petersen

etal.

(1977),

Noble & Silberman (1984),

Deen(1990)

Siddeley & Araneda (1986),

J a n n a s a i ( 1 9 9 0 )

Losada-Calderon & McPhail

(1994), Losada-Calderon

et al.

(1994)

Arribas

etal.

(1995a)

principal geologic environments (Bethke 1984;

Ryeet al.1992): (1) by the disproportionation of

magmatic SOj to H

2

S0

4

and H

2

S following

absorption by groundwater (magmatic-

hydrothermal), (2) by atmospheric oxidation of

H

2

S in the vadose zone over the water table,

associated with fumarolic discharge of vapor

released by deeper boiling fluids (steam-heated),

and (3) by atmospheric oxidation of sulfides

during weathering (supergene). Magmatic-

hydrothermal alunite occurs with minerals such as

diaspore, pyrophyllite, kaolinite, dickite, and

zunyite, which are typical of hypogene (T= 200-

350 C) acidic conditions (advanced argillic

assemblage; Meyer & Hemley 1967). This type of

alunite is characteristic of

HS

deposits, but it may

also appear in areas of advanced argillic alteration

void of ore mineralization

e.g.,

Iwao 1962; Hall

1978). Alunite in steam-heated environments

forms with kaolinite and interlayered illite-

smectite at about 100 to 160 C where fumarolic

vapor condenses above the boiling zone of

neutral-pH,

H

2

S-rich

fluid, typical of geothermal

systems that form low-sulfidation deposits.

423

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A . A rribas, Jr.

Vuggy silica

Quartz alunite

PropylWc

Chlorite-rich

rock

Montmorillonite-rich

rock

100 m

Figu re 2. Cross-section of alteration zones characteristic of h igh-sulfidation . deposits ,.as observed at the

Summitville Au-Cu deposit, Colorado. Diagramat left (simplified from Steven & Ratte 1960) shows schematic

outward zonation from a subvertical mineralized body, shown at right (from Stoffregren 1987).

Because of the relatively shallow and dynamic

environment of mineralization, overprinting

among the three types of acid-sulfate alteration

(including supergene) is possible; however, the

spatial relation of each type of alunite to ore is

different, and correct identification is important

for exploration (Rye

el al.

1992; White &

Hedenquist 1995).

D I S T R I B U T I O N , A G E A N D E C O N O M I C

S I G N I F I C A N C E

In common with other magmatic-

hydrothermal deposits (e.g., porphyry copper

deposits), HS deposits coincide worldwide with

plutonic-volcanic arcs. This association is best

observed in the Cenozoic deposits of the Circum-

Pacific and-the Balkan belt of southeastern Europe

(Fig. 1). These deposits occur in two main

settings: in island arcs and at continental margins.

The tectonic regime during formation of the

deposits seems to be dominantly extensional

(Sillitoe 1993). Some deposits (e.g., Goldfield,

Rodalquilar, Summitville) formed in intra-

continental regions during periods of extension

that followed regional compression and sub-

duction by several m.y.

Tertiary HS deposits predominate, and only a

few deposits are Mesozoic (e.g., Pueblo Viejo,

Zijinshan), Paleozoic (e.g., Temora and others in

southeastern A ustralia), or PreCambrian (the early

Proterozoic Enasen Au deposit located in the

Baltic shield of central Sweden; Fig. 1). The

youngest deposits are Pleistocene (

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O

CM

Figure

3 .

K

2

0 versus Si0

2

variation diagram

for rocks thought to be genetically related to

high-sulfidation deposits. The samples from

12 deposits or districts ( = 140) define a

small compositional field, which contrasts

sharply with the large field defined by

volcanic rocks associated with low-

sulfidation or intrusion-related Au deposits

(>100 samples from 16 districts; Sillitoe

1991b, 1993; Muller & Groves 1993). The

degree of alteration of the rock samples and

precision of the analytical data are largely

unknown; however, according to the

individual data sources, most ofthesamples

are unaltered or very weakly altered. Circles

indicate average values for each high-

sulfidation deposit or district: Ch =

Chinkuashih, Cq = Choquelimpie, Go =

Goldfield, In = El Indio, Ju = Julcani, La =

Laurani, Le = Lepanto, Mo = M otomboto,

Na = Nansatsu, PP = Paradise Peak, Ro = Rodalquilar, Su = Summitville. Compositional fields after Keithet al.

(1991). See Appendix for references and information on data plotted.

50

60

SCv, (wt%)

70

80

similar to that of mineralization. Where abundant

radiometric ages are available, the age of the host

rocks and the age of mineralization are within

analytical precision; where a difference is

indicated, it is typically less than ~1.0 m.y. (Table

2) . A common spatial association exists between

the deposits and shallow, typically porphyritic

intrusions. These intrusions are interpreted to be

the roots of volcanic domes or the feeders of

central-vent volcanoes or maar-diatreme com

plexes, the three main volcanic settings for HS

deposits (Table 2). Some deposits are hosted

entirely within a single dome (Summitville), or

within a dome complex (Julcani). In most cases

the mineralization extends from the subvolcanic

intrusion into country rocks, such as the Main

Vein Cu-Au-Ag deposit and associated breccia

deposits in the Penshan area of the Chinkuashih

district. Some deposits, however, do not show any

(known) spatial association with subvolcanic

intrusions thought to be genetically related to

mineralization

(e.g.,

Nalesbitan, Nansatsu). In the

Rodalquilar Au deposit, dykes and small

intrusions of hornblende andesite which are

interpreted to be temporally related to the

mineralization represent only a fraction of the

altered and mineralized area exposed at the

present depth of erosio n; a larger intrusive body is

thought to exist at depth (Arribas et al. 1995a).

The main control on location of mineralization at

Rodalquilar is the structural margin of two nested,

resurgent calderas. With the exception of

Rodalquilar, the role of calderas in the formation

of HS deposits seems to be limited to facilitating

the emplacement of late intrusive magma along

preexisting caldera ring-fractures (Rytuba et al.

1990).

The magmas thought to be genetically related

to HS deposits have a remarkably limited

compositional variation. The ranges of wt.% K

2

0

and SiC2 for twel ve d epo sits ov erlap g reatly and

show a dominance of calc-alkaline andesitic and

dacitic compositions, with subordinate rhyolite

(Fig. 3). Intermediate calcic volcanic rocks are

limited to porphyritic intrusions in the Lepanto

and Motomboto Cu-Au-Ag districts , and

intermediate-to-felsic alkali-calcic rocks are

characteristic of the Summitville and Laurani

districts (Fig. 3). Interestingly, no deposits have

been discovered in association with alkaline or

mafic magmas, even though these magmas can be

genetically related to low-sulfidation and

intrusion-related Au deposits (Sillitoe 1991b,

1993;

Muller & Groves 1993; Richards this

volume). The data shown in Figure 3 suggest a

relation exists between magma composition and

425

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A . Arribas,Jr.

Table3.Main alteration and m ineralization characteristics of

14

selected high-sulfidation

epithermal deposits

Deposit

Motomboto

Nalesbitan

Lepanto

Chinkuashih

Zijinshan

Nansatsu

Summitville

Goldfield

Paradise Peak

Pueblo Viejo

Julcani

Lateral alteration zoning

(outward from minera

lized bodies)

VS-qtz-alu-*qtz-kao-

kao-sme-Mll-chl

Silicified Hbx-*qtz-kao-

alu-ill-sme-chl-cal

VS/MS-* tz-alu-kao-*

kao-qtz-ill-*chl-ill

VS/MS-*qtz-alu-kao-

ill-chl-kao

VS/MS-*qtz-dic-alu-qtz-

dic-ser-qtz-scr

VS/M S-* alu-dic-pyo-*

ill-kao-sme-*

PRO

VS(MSh*qtz-alu--

qtz-kao-*kao-iU-*

sme-chl

MS{VS)-*qtz-alu-kao-*

iII-sme-

PRO

Vertical (due to deposit

style): MS(VS)-*

qtz-alu-kao- sme-chl

Complex + overprinted

Pre-ore:VS/MS-qrz-alu-

Venical alteration

zoning

(shallow to deep)

VS/MS-qtz-alu-*qtz-

kao-*ill-kao-*cbl

Silic if ied Hbx-*qtz-kao-

alu- ill-sme-chl-cal

M S A ' S - * A A - S E R -

(K-silicate in subjacent

FSE porphyry copper)

VS/MS-* qtz-dic-lu- qtz-

dic-ser-qtz-ser

VS/MS-*alu-*dic - ser -

py-ser -chl -PRO

VS(MS)-qtz-kao-

alu-*qtz-kao-*SER

MS>VS-*qtz -a lu-

kao-*qtz-kaopyo

MS(VSHqtz -a lu-kao

(SER in faulted, deeperC?)

East Zone deposit)

Early: Kao-py-qtz-*

alu-py-qtz

Late: MS-pyo-dia

Princial ore minerals

Py, ena-luz, mar, sph, gal, len-

tet, are, cpy, arg, nat.Au, tell

Py, chalc.qtz, ceo, bor, cov,

ena, tell

Ena-luz, py, ten-let, cpy, py, ele.

sph, gal, mar, sele, tell, Sn-

bearing sulf

py, ena-luz, fam, ten-let,nat.Au,

ele ,

bar, naLHg, tell, sph, gal,

cpy, geo, bou

py, dig, ena, cov, mol, naLAu

cpy, bor, let-ten, gal, sph

ena-luz. py, ele, nat.Au, arg,

pyr, cpy, bor, sph, gal, cas, sta,

mol, can

py, ena-luz, cov, mar, naLS,

nat.Au , sph, gal, bar, cpy, ten

py, fam, ten-let, bis, got,

naLAu, ena-luz, bar, tell, sph,

cov

bar, stb, bis, nal.Au, mar, py,

nat.S, cin, sph, gal, cpy, ars,

let, arg, cov, fam

py, sph, ena, nal.Au, nat.S, bar,

len-tet, fam, gal, bar, stb, ele,

sele, tetl, Bi- Pb- Ag- sulf

py, wol, cas, nat.Au, ena, luz,

Ore

mincratizaUon

in:

Silica core

Silica core

Silica core

Silica core

Silica core

Silica core

Silica core

Silica core

Silica core

I n A A +

MS zones

Veins

Ag/Au

3 5 - 4 5

Very

low Ag

4

2

N/A

8/11/2019 1995 Arribas

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Table

3

(continued)

High sulfidation Epithermal Deposits

Deposit

Motomboto

Nalesbitan

Lepanto

Chinkuashih

Zijinshan

Nansatsu

Summitville

Goldfield

Paradise Peak

Pueblo Viejo

Julcani

El Indio

LaMejicana,

Nevadosdel

Famatina

Rodalquilar

N and order of

main mineral

events

2

AA-> sulfide-Au

2

Py-qtz-Cu-Au

3

AA-CuAu

-Au

2

AA-Cu-Au +

late bar-Au

SeveraK?),

Qtz-ser-dic-*qtz-

alu-dic- sulfide

2

AA-*Cu-Au

3

AA-CuAu-+

bar-base-metal-Au

5+

AA-ena-Au-

ten-py-Bi-Tc-

3

AA-Au-Ag-+

Hg

2

Alu-kao-py-Au-

MS-pyo-dia-Au

Several

AA+VS-tou

breccias-* sulfide

veins (main, late)

2

AA+Cu- Au

2

AA-silica-py-Au

Inferred

mineralization

depth (m)

1

Unknown

300-500

flinc

Unknown

500

Unknown

150-300

flinc

400-500

flinc + geol

100-300+

geol

Unknown

Shallow; lacus

trine sediments

preserved

200-300

200-300

flinc + geol

Inferredore-

forming

mechanism

Unknown

Boiling (Hbx )

2

Mixing/cooling

Unknown

Mixing/boiling

Mixing/cooling

Cu-Au by

mixing, bar-Au

by oxidation

Mixing/cooling;

oxidation

Boiling, with

Hbx in early Au-

Ag stage

Sulfidation +

boiling

Mixing + boiling

Mixing/cooling

Boiling (Hbx)

2

+

mixing/cooling

Supergene ox idation/

secondary Au

enrichment?

Irregular to 100 m

Complete to 130

m;

yes

Not important

Important in upper

250nv, yes

Widespread to 100

nv,

yesbut may

besteam-heated

Irregularto100 m;

may be partly

steam-heated

Widespreadto80

nv, supergene

alunite(-lOMa)

Widespread

to

250

m;supergene alu-

nite (100.5 Ma)

Widespread to

- 100 m

Widespread

to

80

m;supergene

alunite (4-3 Ma)

References

Perello(1994)

Sillitoe e lal.(1990)

Garcia (1991), Claveria

Hedenquist(1994)

Huang (19 55),

Tan el

al.

(1993)

Ren el

al.

(1992),

Zhangetal.(1994)

Izawa

Cunningham (1989),

Hedenquisl

elal.

(1994a)

Steven & Ratte (I9 60,)

Stoffregen (1987), Gray &

Coolbaugh(1994)

Ransome(1907, 1909).

Ashley (1974), Vikre (1989,

written comm., 1995)

Johneial. (1991),

Sillitoe & Lorson (1994)

K es l era i ( 1981)

Russell AKesler (1991).

Munteane/a/. (1990)

Deen (1990), Rye (1993)

Siddeley Araneda (1986),

Jannasrtai(1990)

Losada-Calderon McPbail.

(1994 ), Brodtkorb

Paar

(1993)

SSInger-von Oepen

et

al.

(1989)Arribasfl/.

(1995a)

development of the oxidized and reactive

magmatic vapor plume that is thought crucial to

the formation of HS deposits.

DEPOSIT FORMAffDCONTROL:

CLASSIFICATIONS

High-sulfidation deposits display a wide

variety of styles of mineralization that includes

veins,

hydrothermal breccia bodies, stockworks,

and disseminations or replacements. This

variation in the structure of the orebodies is

complemented with variations in other deposit

features, including ore and alteration mineralogy,

paragenesis, and metal ratios (Tables 2, 3). In

addition, some deposits present complex relations

which may be composite,e.g.,between high- and

low-sulfidation mineralization

e.g.,

quartz-Au-

stage veins at El Indio, and some of the veins at

Julcani). Definition of styles ofHSmineralization

427

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8/11/2019 1995 Arribas

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h > B

Lepanto fault

Age (Ma)

Ore

deposits and lithology

1.56-1.17

I [ High suttkJation Cu-Au ore

1.45-1.22

E%3 Porphyry Cu-Au ore

1.2-0.9 ( i l l P ost-mineralization cover

t 3 Ouartz-oloiUe porphyry

22-1 .8 r I M Dacte porphyry & pyroclastics

Cre.-Mto. B Basement rocks "

m

Jea level

raaftHi

)n/mineralizatJon

m

Stratabound f^,

-700m

JSWfl

pgg Vuggy silica/

BSa massive silica

RS|9 Advanced

^ ^ argiHic

f = g ArgiHic

1 Propyiitic

Figure 4. Longitudinal (A) and transverse (B) cross-sections of the Lepanto-FSE Cu-Au-Ag deposits (Philippines),

showing structural and Uthologic controls on formation of the high-sulfidation and porphyry-type ores (simplified

from Garcia 1991). Potassium-argon dating of country rocks and alteration m inerals associated with the porphyry and

high-sulfidation deposits indicates that hydrothermal Cu-Au mineralization took place in the middle of

a

Pliocene to

Pleistocene event of dacitic-andesitic magmatism (Arribas et al. 1995b). Note the overall spatial overlap of the

magmatic and hydrothermal "plumbing" systems (i,e., volcanic vents of Pliocene dacite, quartz diorite intrusions,

porphyry deposit, and deeper parts of epithermal mineralization).

The zones of alteration with increasing depth

typically grade from a shallow silicic zone

through advanced argillic, argillic, argiHic/

sericitic, into a sericitic or phyllic zone with

quartz, sericite, and pyrite. This alteration

sequence occurs over a vertical interval that

ranges from a few hundred meters to more than

1000 m, and has been best documented by deep

drillholes in the deposits of smaller size, in which

the vertical span of mineralization is less than

about 300 m (e.g., Rodalquilar, Summitville; Fig.

5B). At Lepanto, sericitic alteration at depths of

400 to 500 m below the epithermal deposit gives

way, laterally towards the south, to K-silicate

alteration of the FSE porphyry Cu-Au deposit.

Porphyry-type stockwork mineralization at

Paradise Peak is contained within the sericitic ores

of the East Zone deposit which, according to

Sillitoe & Lorson (1994), formed underneath the

main HS orebodies in the area. A quartz-sericite-

pyrite zone with trace am ounts of chalcopyrite and

molybdenite surrounds an intrusion of monzonite

porphyry 300 m below the HS deposit at

Summitville (Gray & Coolbaugh 1994).

The lateral and vertical alteration zones

described above correspond to a generalized

model. They are useful in exploration because

they help in understanding the genetic environ

ment of a deposit and provide spatial "markers"

within the extinct hydrothermal system.

Experimental data on the relative stability of

minerals such as alunite, kaolinite, pyrophyllite,

and diaspore (Hemley et al. 1969, 1980), coupled

with the temperature ranges noted for these and

other related acid minerals in active systems

(Reyes 1990; Reyes

et al.

1993), also provide

information that contributes to definition of the

paleoconduits in extinct systems.

If studied in detailed, several superimposed

and crosscutting stages of pervasive as well as

fracture (conduit)-related mineralization may be

recognized in the majority of deposits. These are

the expected result of variations, during the course

of mineralization, in temperature, pressure, and

composition of the hydrothermal fluid and the

degree of wallrock interaction. Detailed field and

petrographic studies at the Monte Negro orebody

in the Pueblo Viejo deposit have resulted in

429

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A. A rribas, Jr.

Rodalqultar caldg--r?

m r i n ^ < ^ ^ ' '

^KmSmm

iominacakieraV ftt ls / Un

s, margin

N Mm

rffflsRSm

^ ^ V ^ \ jHodafottf/ar

i Vuggy sil ica " ^ -

||l ,

r = j Advanced argilltc V. _ i p ^ l t t j r

TO3Argillic ^^^5i2Sii5S>

liS

*

i

^^^^^\.._ f;||flp-:

:

f

L

A

I Serkatic

p l l Propylitic

GOOD In tens e sup er ge ne a ctd-sulfate overprint

40 0.

0,

-400

Elevation

Lower limit of

sulfide oxidation

l i lHi

A JCSJSSw

A

A

vf*

A A A

, A A A

[A A A

Au-{Cu-TeSn) Ngh-

sutfidation deposits

:ij|HrAI

% A A X

A A A /

ft. A A X.

A A A J

* A 4* ,

A A A /

Pb-Zn-{Cu-/

quartz vein

r

s

Figure 5. Generalized surface alteration map (A) and cross-section (B) of

the

Rodalquilar

HS deposit in the Rodalquilar and Lomilla calderas, southeastern Spain (from Arribas et

al .1995a). The boundaries shown between alteration zones are irregular and gradational.

identification oftwo stagesofmineralization,

interpreted to correspond to two distinct magmatic

pulses (Muntean

et

al. 1990). During the first

stage (responsible for ~60% of the Auinthe

deposit), shallow kaolinite-quartz-pyrite and deep

alunite-quartz-pyrite-quartz zones were

de-

veloped, with gold mineralization in association

with disseminated pyrite in the wallrock; during

the second stage (responsible for about 40% of the

Au), an extensive zone of siliciflcation with pyrite

sphalerite enargite veins formed at shallow

levels, aboveazoneof pyrophyllite-diaspore

alteration (Munteanet al. 1990).

O R E A ND G A N G U E M I N E R A L O G Y , A ND

T I M I N G O F M I N E R A L I Z A T I O N

White et al. (1995) and White & Hedenquist

(1995) presented detailed discussions on various

aspects of epithermal gold mineralization on the

basis of observations from alarge number of

deposits around the Pacific; their conclusions with

respecttoore and gangue mineralogy inHS

deposits are included here, inaddition tothe

particular features of the deposits listed in Table

3.

Pyrite and enargite (and its low-temperature

dimorph luzonite) are the dominant sulfides in HS

deposits; pyrite

is

abundant but the amount of

enargite and luzoniteisvariable. Common ore

minerals, listed by decreasing abundance from

variable to very minor, include tennantite-

tetrahedrite, covellite, native gold and argentian

gold (electrum), marcasite, chalcopyrite, spha

lerite, and galena. Famatinite is locally abundant

in some deposits (Goldfield, La Mejicana). Sparse

ore minerals include bornite, cassiterite, cinnabar,

molybdenite, orpiment, realgar, stibnite, and

wolframite (the last locally important at Julcani).

Other minerals present inminor amountsin

several deposits include Pb-, Ag-Pb, Bi- and Sn-

bearing sulfosalts (Table 3).

Fine-grained quartz is the dominant gangue in

HS deposits. Other common but minor gangue

minerals include barite, kaolinite, alunite,

pyrophyllite, diaspore, and Ca-,Sr-, Pb- and REE-

bearing phosphate-sulfate mineral(s) suchas

svanbergite-woodhouseite orcrandallite

(Stoff-

regen & Alpers 1987). For example, high-grade

430

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vein specimens from Chinkuashih, Goldfield, and

La Mejicana have spectacular intergrowths of ore

minerals with kaolinite, alunite, or pyrophyllite.

This observation implies that ore formation

occurred under modera te ly ac idic to ac idic

conditions, which are inconsistent with transport

of Au as Au(HS)2" (Seward 1973) . Recent

studies of Au solubility in high-temperature acid

sulfide solutions have resulted in identification of

AuHS as one of the principal gold complexes in

HS mineralization (Bening & Seward 1994), the

other possibility being AuCl2~(e.g., Hedenquist

et al. 1994a).

The number and order of mineralizing events

provide critical information for reconstruction of

the hydrothermal system that results in HS

mineralization. A minimum of two stages of

alteration/mineralization has been recognized in

most deposits on the basis of crosscutting

relations (Table 3). The most common evolution

is from an early leaching and alteration stage to a

later ore-forming stage. Vuggy silica rock and the

advanced argillic assemblage with disseminated

pyrite form typically early-stage acidic alteration,

and are followed by Cu Au Ag deposition.

Detailed studies in some districts (e.g., El Indio,

Lepanto), however, have resulted in identification

of two metal stages, an early Cu-rich, Au-poor

stage, dominated by enargite-luzonite, and a late

Au-rich, Cu-poor stage, associated with

intermediate-sulfidation-state sulfides such as

tennantite-tetrahedrite and chalcopyrite, and

tellurides. The transition from quartz-alunite-

pyrite alteration to enargite-pyrite and finally to

tennantite-tetrahedrite, the last typically without

sulfate (alunite) but with quartz-sericite gangue

and wallrock alteration, indicates a fluid

progressively more reduced and less acid. At

Summitville and Chinkuashih (also Tambo and

Furtei-Serrenti; Table 1), a late stage of barite-

gold has been documented.

C H A R A C T E R I S T I C S A N D S O U R C E S O F

H Y D R O T H E R M A L F L U I D S

Results of recent detailed fluid-inclusion and

stable-isotopic studies reveal much about the

composition, temperature and sources of

hydrothermal fluids in HS deposits. Combination

of these data with geological and mineralogical

observations mentioned above allows the nature

of the altering and ore-forming fluids to be

determined. The framework for the interpretation

has benefited from information on the compo

sition and fluxes of volcanic discharges and active

magmatic-hydrothermal systems (Hedenquist &

Lowenstern 1994; Giggenbach this volume;

Hedenquist this volume).

Fluid-inclusion Evidence

Suitable hosts for fluid-inclusion studies are

scarce in HS deposits, as the gangue minerals are

typically fine-grained and even millimeter-size

hydrothermal quartz crystals are usually late stage

and vug-filling. Satisfactory results are obtained

on secondary fluid-inclusions in igneous quartz

phenocrysts from altered wallrocks; although

lacking temporal information, these inclusions

seem to provide a representative cross-section of

the fluids involved. The most reliable data on the

ore-forming fluids are obtained through infrared

microscopy directly on ore minerals, such as

enargite (Deen 1990; Mancano & Campbell

1995).

The temperatures and salinities estimated for

HS deposits define a wide range, from 90 to 480

C and 300 C)

fluids of variable salinity, which have been

documented in several deposits and are generally

431

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A .

Arribas, Jr.

Table 4. Summary of fluid-inclusion microthermometric data for high-sulfidation deposits

Deposit

Motom boto, Indonesia

Nalesbi tan, P hil ippines

Lepanto, Phil ippines

Chinkuashih, Taiwan

Zijinshan, China

Nansatsu. Japan

.^

Akaiwa, Japan

Mitsumori-Nukcishi , Japan

Summitvi l le, Colorado

Goldfield, Nevada

Paradise Peak, Nevada

Julcani, Peru

Ccarhuaraso, Peru

Colquijirca, Peru

Can-Can (La Coipa) ,

Chile

El Indio, Chile

La Mejicana (LM) and

Nevados Famatina (NF),

Argentina

Rodalquilar, Spain

. .

Funei-Serrenti, Italy

Host-mineral

studied

Barite

Quartz

Enargite

Quartz, barite,

al unite

Quartz

(no

details

reported)

Quartz

Diaspore

Quartz, barite,

quartz phenoe

Quartz phenoe

Barite

Quartz phenoe

Quartz, barite

Quartz, barite

Quartz

Quartz phenoe

Quartz phenoe

Wol, ena, quartz

Siderite

Quartz phenoe

Quartz phenoe

Sphalerite, quartz

hubnerite

Quartz phenoe

N/A

Quartz, quartz

phenoe

Quartz, barite,

quartz phenoe

Temperature

(C)'

150-180

220-260

170-290

180-330

160-300

220-380

100-160

(300-420)

130-250

-270

250-310

190-240

210-330

180-280

(300-390)

-100

230-480+

210-280

(370-410)

180-210

300-380

(up to 450)

160-280

360-450

230-330

220-250

330-380

230-260

170-350

190-280

140-180

(>300)

200-460

160-340

230-480

170-300

220-450

190-320

90-140

(390-500)

Salinity

Associated

(equiv wt.9E- NaCI) alteration

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Table4.(continued)


Deposit

Comments References

Motomboto, Indonesia

Nalesbitan, P hilippines

Lepanto, Philippines

Chinkuashih, Taiwan

Zijinshan, China

Nansatsu, Japan

Akaiwa, Japan

Mitsumori-Nukcishi, Japan

Summ itville. Colorado -

Goldfield, Nevada

Paradise Peak, Nevada

Julcani, Peru

Ccarhuaraso, Peru

Colquijirca, Peru

Can-Can (La Coipa),

Chile

El Indio, Chile

La Mejicana (LM) and

Nevados Famatina (NF ).

Argentina

Rodalquilar, Spain

Furtci-Serrenti, Italy

Reconnaisance study in late-stage barite

Reconnaissance study; liquid

CO2

observed

Sampled interval 3 km long by 0.5 km high ; cooling fluids

away from subjacent porphyry Cu-Au deposit, where

T

h

>450C & salinity up to 54 eq wt.% NaCl

Poorly-documented samples along a 450-m vertical interval;

the higher T|,s in samples at -750 m depth; CO2 observed

Associated with main stage Cu

Deep alteration zone (>600mdepth)

Associated with late, shallow silica-Au

Associated with early silica and quartz-dickite

Late,vug-filling quartz

Qtz in breccia, saline liquid and low-salinity vapor coexist

Vein quartz -40 0mbelow Kasuga deposit

Coarse-grained diasporc

Not (known) Au or Cu mineralization, but high salinity

fluids

Liquid-rich; salinity >6 eq

w..1o

NaCl only in vuggy silica

associated w ith Cu mineralization;CO2observed

Liquid- and vapor-rich inclusions; also polyphase inclusions

Late barite-Au assemblage

True T

h

is interpreted to be 250-290C

Hydrostatic and near-lithostatic pressures suggested

Late,

vug-filling crystals in hydrotherma breccia;

From stockwork Au East Zone deposit; CO2 observed

Quartz-alunitepyrite

Pre-ore tourmaline breccia dykes, lithostatic pressures likely.

Main-stage ore fluids, also inner veins, liquid-rich inclusions

Late-stage ore fluids, also in outer ve ins; P correction applied



Two generations identified; both may be very saline. Evidence

for P above hydrostatic and higher salinities at depth

Copper and gold stages

Late stage

Interpreted as early, with vapor-rich inclusions, CO2 observed

LM &

NF:

includes liquid-, vapor-rich and polyphase inclusions

NF:

complete transition from porphyry-type fluids in K-

silicate stage (SOO -oW^C, up to 67 eq wt% NaCl)

through sercitic to epithermal fluids in HS (AA ) stage;

vapor-rich inclusions typically less saline

Vertical temperature and salinity gradient: high-temperature

brines coexist with low -aiinity vapor inclusions;

hydrostatic and near-lithostatic pressures suggested

Includes high + low-salinity fluids(22-23,

8/11/2019 1995 Arribas

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A .A rribas, Jr.

and Julcani (Deen 1990) are broadly similar, but

their salinities are distinctly different (0.2-4.5

equiv. wt.% NaCl versus 8-18 equiv. wt.% NaCl,

respectively), providing constraints on the role of

a saline magmatic liquid (versus low-salinity

vapor) in the generation ofHSdeposits.

Group 3. Lower temperature (e.g., 90-180

C), dilute (typically 600-m vertical

interval (extending 500 m below the ore zone; Fig.

6) shows a gradient which correlates with the

change in dominant alteration, from silicic and

advanced argillic (T=170-300 C, salinity = 2-15

equiv. wt.% NaCl at the elevation of the orebody)

to sericitic (7"= 220-450 C, salinity = 2-45 equiv.

wt.% NaCl) assemblages.

The transition from advanced argillic alteration,

through quartz-sericite-pyrite, to K-silicate

alteration and typical porphyry-type high-

temperature (600+ C) and high-salinity (up to 67

equiv. wt.% NaCl) fluids of magmatic origin is

displayed, among the examples reviewed, at the

Lepanto-FSE and La Mejicana-Nevados del

Famatina epithermal-porphyry copper systems.

The cooler and less saline inclusion fluids

documented in the ore zone of the HS deposits are

interpreted to reflect mixing of magmatic and

meteoric fluids in an environment shallower than

that of porphyry mineralization. Furthermore, in

common with porphyry-type deposits, high-

temperature, vapor-rich, low-salinity fluid

inclusions coexist with high-temperature, liquid-

Temperature (C)

200 300 400 500

400 -

Elevationof

Ontodeposits

200

3

I

1

I

-200

-400

HjO +

5

wt% NaCl

(hydrostatic)

(hydrostatic)

(lithostatic)

B

1

s -|400

>

3

I

200

- 600

1

I

.

o

800

Figure 6. Elevation versus temperature diagram

showing the range (horizontal line) and average

(vertical line) of fluid-inclusion homogenization

temperatures measured in the Rodalquilar Au deposit,

Spain. Also shown are the temperatures calculated, on

the basis of8

34

S

su

ir,d

M

uifatt

for four

coexisting alunite-

pyrite samples (large filled circles), reference boiling-

point curves, and vertical spans ofthe alteration zones

mentioned in the text. Estimated salinities of fluid

inclusions in the shallow advanced argillic/silicic zone

and deep sericitic zone range between 2 to 30 equiv.

wt.% NaCl and 2 to 45 equiv. wt.% NaCl, respectively

(modified from Arribasetal.1995a).

rich hypersaline inclusions (i.e., with Groups 1

and 4, above). These fluids may be the result of

boiling of a high-temperature liquid, or they may

reflect immiscible vapor and hypersaline liquid

derived directly from shallow-emplaced magma

(Rye 1993; Hedenquist & Lowenstern 1994;

Shinohara 1994; Hedenquist this volume).

Sulfur-isotope Evidence

The abundance of coexisting hydrothermal

sulfides and sulfates, in addition to the possibility

434

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.Sulfides Sulfates T

V -

5

3 4

SJS

Lepanto

Chinkuashlh

Nansatsu

Summltvil le

Goldfield

Pueblo Viejo

Julcani

El Indto

Rodalqullar

+

+

+

i

T

I

I

- H 1 1 1 1 1 h-

-10 0 10 20

*S

(%., CDT)

A

3 4

S H

2

S -S O 4

Temp . (C)*

220 - 420

220 - 270

200 - 240

2 0 0 - 3 9 0

200 - 350

1 8 0 - 2 6 0

2 1 0 - 2 7 0

H

2

S / S 0

4

2-6

-

3

4

-

-

5

220-330 5

'(mineral pairs)

30

Figure 7. Range of 8

34

S (per mil) values for sulfides and sulfates from nine high-

sulfidation deposits. Also shown are the values calculated for8

34

S for total sulfur in the

hydrothermal system (triangles), H

2

S/S0

4

, and the range of temperatures determined

from sulfide-sulfate mineral pairs. Solid triangles indicate deposits in which 8

34

S

S

was

calculated on the basis of isotopic analyses of samples of unaltered whole rock

genetically related to mineralization. See Appendix for references and information on

data plotted.

of measuring

3 4

S / S in host rock and genetically

related igneous rock (Sasaki et al. 1979), allows

sulfur-isotope studies to provide information on

the composition, temperature, and sulfur sources

of the hydrothermal fluids. The results of detailed

studies in nine HS districts show a remarkable

consistency (Fig. 7). In agreement with the

observations in active volcanic-hydrothermal

systems (e.g., Kiyosu & Kurahashi 1983), sulfide

and sulfate minerals are mainly in isotopic

equilibriu m, and, therefore , their overall S/ S

depends on the temperature of mineralization and

th e

34

S/

32

S of total sulfur in the hydrothermal

system. Only the data for alunite from the

Campana vein in El Indio (Fig. 7) are different. If

the measured El Indio alunites are not steam-

heated or supergene (unlikely as they contain fine

grained pyrite; Jannaset al. 1990), the most likely

explanation is a "magmatic-steam" (Rye et al.

1992) origin, in which the8

34

S of alunite is close

to the composition of total sulfur in the system

(e.g., Alunite Ridge in Marysvale; Cunningham et

al . 1984; Rye et al. 1992) . Com bined with the

8 S values of pyrite and enargite from the sam e

vein, these values indicate drastic changes in

H

2

S/S0

4

during the course of mineralization

(similar to those for the Red Mountain alunite

deposit; Boveet al. 1990; Rye 1993).

The main conclusions of the sulfur-isotope

studies in HS deposits are: (1) sulfur in the

deposits is magmatic, but the magmatic sulfur is

overall heavier than mantle values (from

5

34

S

= 2

0 / 0 /

2 'oo at Sum mit ville, to 9 2 'oo at R odalquila r;

Fig. 7). This is not surprising given the most

common geological setting of the deposits;

isotopically heavy igneous sulfur is common in

volcanic arc environments (e.g., Ueda & Sakai

1984). (2) A simple mass-balance calculationdone in several deposits using the S/ S values

of the igneous rocks and the average

34

S/

32

S

values of sulfides and sulfates indicates that

H

2

S/S0

4

in the hydrothermal fluids was generally

about 42 (Fig. 7; Ryeet al. 1992; Hedenquist et

al . 1994a; Arribas et al. 1995a). This is a

minimum value for ore-forming fluids because it

applies mainly to the early stage of hydrothermal

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8/11/2019 1995 Arribas

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A . Arribas, Jr.

alteration, which is characterized by a sulfate-rich

alunite-pyrite assemblage. (3) Isotopic equilib

rium between sulfide and sulfate in the

hydrothermal solutions results, in a majority of the

deposits, in reliable temp eratures calculated on the

basis on A S

H2

s-so4 (Fig- 7). Pyrite-alunite

mineral pairs were used most commonly, and

where sam pling with depth is available, they show

a thermal gradient:

e.g.,

220 to 330 C over 200-m

elevation at Rodalquilar (Arribas

et al.

1995a),

200 to 390 C over -900 m at Summitville (Rye

1993); 220 to 420 C over 500 m at Lepanto

(Hedenquist and Garcia 1990; J.W. Hedenquist,

unpub. data). Other mineral pairs used with

consistent results include pyrite-barite (Vikre

1989; Deen 1990), sphalerite-barite (Vennemann

et al.

1993), and pyrite-gypsum (Vikre 1989). The

range of isotopic temperatures is consistent with

temperatures estimated from fluid inclusions and

alteration mineralogy (e.g., Hemley

et al.

1980;

Reyes 1990; Reyes

et al.

1993). The range is also

consistent with formation of alunite at

temperatures below ~400 C, when S0

2

gas starts

to disproportionate in the hydrothermal solution

(Sakai & Matsubaya 1977; Bethke 1984).

Oxygen- and Hydrogen-isotope Evidence

In terms of oxygen and hydrogen isotopic

composition, the fluids that form HS deposits are

arguably some of the better documented and

understood in ore-deposit studies. This situation

contrasts sharply with that of a decade ago, at

which time no d ata were av ailable to corroborate

the affinity-suggested between fluids in active

volcanic-hydrothermal systems and HS deposits

(e.g., Heald

et al.

1987; Hedenquist 1987). Stable-

isotope studies of HS deposits are particularly

illuminating because of: (1) the abundance and

variety of oxygen- and hydrogen-bearing minerals

(e.g., alunite, illite, kaolinite), (2) the developm ent

of analytical procedures for complete stable-

isotope analysis of alunite, including 8

l 8

O

s o 4

and

6

1

0

O H

that help to distinguish the various types

of alunite and associated acid-sulfate alteration

(Rye

et al.

1992; Wasserman

et al.

1992), (3)

fewer limitations on the interpretation of the

isotopic data because of the relatively young age

of mineralization of most HS deposits and general

lack of post-depositional effects that disturb the

stable-isotope systematics, and (4) the availability

of detailed information on the isotopic

composition of fluids in active geothermal and

volcanic-hydrothermal systems, which allows

fluids estimated in HS deposits to be compared

with those in their active equivalents.

Some limitations still exist. These may be

independent of obvious factors such as sampling

or mineral-preparation procedures (fundamental

for achieving representative and reliable results),

analytical imprecision, and natural variations, as

observed in active sy stems (e.g., Aoki 1991, 1992;

Rowe 1994). Important limitations that must be

taken into account for optimum use of the stable-

isotope data are related to (1) the choice of

temperature of mineral formation for calculation

of the fluid isotopic composition, (2) the lack of

mineral-water fractionation factors for some

minerals (e.g., pyrophyllite), and (3) the

disagreement among fractionation constants

proposed for even common minerals such as illite

(see Dilles

et al.

1992, for a discussion) and

kaolinite. For example, at 200 C there is a

differenc e of 20 Aw betw een the D/H frac

tionation constants for kaolinite - water as given

by Marumo

et al.

(1980) on the basis of samples

of minerals and water from active systems, and by

Liu & Epstein (1984) on the basis of experimental

results. For these reasons, discussion of the

sources of water during acidic alteration in the

deposits considered here is based on the average

of the data collected for alunite, for which

fractionation factors are well-know n (Stoffregen

et al.

1994). The magmatic-hydrothermal alunite

typical of-HS deposits gives good results because

it is relatively coarse-grained (post-mineral D-H

exchange is not a problem; Stoffregen

et al.

1994)

and commonly is closely associated with ore, thus

recording equilibrium conditions of a fluid closer

in composition to the ascending mineralizing

solution than the kaolinite or illite from outer

alteration zon es.

Oxygen and hydrogen isotopic compositions

of water in HS deposits are clearly consistent with

mixing between a high-temperature magmatic

fluid of 8

1 8

0 = 9 l/oo and 8D = -30 20^oo and

meteoric groundwaters (Fig. 8). In part because of

4

->

r

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0 -

-2 0 -

-4 0 -

G Alunite alteration stg.

O Ore mineralization stg.

Q Alteration/

mineralization

Subduction-related

volcanic vapor

Arc + crustal

felsic magmas

Acidic fluids In h igh-

] su l f lda t ion depos i ts

.20

W W o

so

6D( .)

-100

i 1 1 1 ' r

-20 -15 -10 -5 0 5

6

1 8

0 (%o, S M O W )

ctive systems

(Giggenbach, 1992b)

Volcanic

Geothermal

10

15

20

Figure 8. Summary diagram showing variation in oxygen- and hydrogen-isotope composition of hydrothermal

fluids in high-sulfidation deposits. The average isotopic composition for the main stages of acidic alteration

(squares) and ore-mineralization (circles) fluids are shown. Where possible, only alunite data were used for the

alteration stage (SD and 8

l8

O

SOi

,); &

I8

0

0H

is not used because hydroxyl oxygen requilibrates with the hydrothermal

fluid during cooling (Ryee tal.1992). Tie-lines between data points connect samples from the same deposit. Inset

shows the isotopic composition of fields defined by waters from active geothermal systems and high-temperature

fumarole condensates in subduction-related andesitic volcanoes (from Giggenbach 1992b). Go = Goldfield, Ju =

Julcani, Le= Lepanto, Nansatsu district: Ka = Kasuga, Iw = Iwato, NF = Nevados del Famatina, PV = Pueblo

Veijo, Ro = Rodalquilar, RM = Red Mountain, Lake City, Colorado, Su = Summitville. The approximate

compositions of groundwaters suggested for several deposits are indicated by the intials parallel to the meteoric

water

line.

See Appendix for references and information on data plotted.

the very light isotopic composition of local

meteoric water, this meteoric-magmatic water-

mixing trend is displayed particularly well by the

three stages of alteration/mineralization at Julcani

(Deen 1990; Rye 1993): from a magmatic-water-

dominated early stage of (alunite) acid-sulfate

alteration (Ju; Fig. 8), through main ore-stage

fluid-inclusion waters (Ju

t

and JU2), to m eteoric-

water-dominated late ore-stage fluid-inclusion

waters (Ju

3

). In addition to Julcani, the ore fluids

at Summitville (Rye et al 1990; Rye 1993) and

Rodalquilar (Arribas

et al.

1995a) also have lower

5

1 8

0 values than those of acidic alteration fluids,

indicating greater dilution by groundwater (Fig.

8) . The extent of an O-shift in the groundwater

component due to water-rock interaction, as

typically seen in some neutral-pH geothermal

systems, is not known, but such a shift is not

indicated by the Julcani data.

The overall oxygen- and hydrogen-isotope

relations are identical to those of volcanic-

hydrothermal and geothermal systems associated

with subduction-related volcanism (Giggenbach

1992b; Fig. 8, inset). The similarity is even closer

between the composition of acidic alteration fluids

(large shaded field, Fig. 8) and the vapor

condensates from high-temperature fiimaroles of

andesitic volcanoes (dark shaded field, Fig. 8,

inset),

such as Nevado del Ruiz, Satsuma

Iwojima, or White Island, the last documented to

have a geochemical en vironment similar to that of

HS mineralization (Hedenquistet al. 1993).

The origin of the D-enriched magmatic (end-

member) fluid of

HS

deposits has been interpreted

in two ways. Most workers conclude that the

acidic fluid in HS deposits is derived from

absorption of magmatic vapors outgassing from

arc volcanoes or felsic magmas in crustal settings

{e.g., Hedenquist & Aoki 1991; Matsuhisa 1992;

Giggenbach 1992a; Vennemann et al. 1993;

437

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(commonly

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A .A rribas, Jr.

ALTERATION

ORE DEPOSITION

B

2

Heated gro und- ..

waters

*

Magmatic vapors

(incl.,

S 0

2

. Ha )

Magmatic

brine

f

Vuggy silica

Alunite

Kaolinite

Sericite

K-silicate

Absorption of 7 *

high-P vapor j

Heated

groundwater

Magmatic

brine

Possible

Cu{Au)

Figure 9. Model showing the two main stages of evolution of HS deposits. A: Early stage of advanced argillic

alteration dominated by magmatic vapor. B, and B

2

: Two genetic hypotheses proposed for the stage of ore

formation. B, = absorption of high-pressure vapor by entrainment in meteoric water cell at depth to explain low-

salinity, mixed magmatic-meteoric ore fluid (Hedenquist this volume). B

2

= ascending metal-bearing magmatic

brine with shallow cooler meteoric waters to explain high-salinity, mixed magmatic-meteoric ore fluid (White

1991;Rye 1993;Hedenquiste tal. 1994a).

metals strongly partitioned into the high-density

liquid (Hemley et al 1992; Hedenquist this

volume).

At this early intrusive stage, several modes of

magma degassing may occur which will lead to

different styles of magmatic-hydrothermal

systems with or without associated mineralization

(Giggenbach 1992a). To form the styles of

alteration and the spatial distribution of alteration

zones characteristic of HS deposits, degassing

must be very efficient, with oxidized high-

temperature magmatic vapor reaching shallow

depths with little reaction with rock or dilution by

groundw aters at g reater depths (Fig. 9A). D ilution

with groundwaters is unlikely because the high

temperatures surrounding the cooling magma

cause meteoric water cells to be displaced from

the magma core (Fig. 9A). In addition to the

relatively low pressure at the depth of intrusion,

effective degassing will be favored by the

structural factors characteristic of HS deposits,

such as fractured volcanic domes or roots of

domes, caldera or diatreme faults, volcanic vent

contacts, and active faults with a dilational

component.

As the high-temperature magmatic vapor

reaches shallow depths of less than a kilometer, it

may be absorbed by groundwater if it does not

discharge as a fumarole. The acidity of this

groundwater-absorbed vapor condensate increases

as the liquid cools, first at temperatures below

~400 C by disproport ionate of S0

2

to form

H

2

S 0

4

and H

2

S (Day & Allen 1925; Sakai &

Matsubaya 1977), then by progressive disso

ciation of H

2

S 0

4

and HCl at lower temperatures

(

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constituting a relatively small part of the mixture

(generally

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A .A rribas, Jr.

vapor is required for transport of sufficient

amounts of metals (Hedenquist this volume;

Sillitoe this volume). These conditions are

consistent with the low salinity of the Lepanto and

El Indio fluid-inclusion data. Mineral deposition

in this case may be caused by mixing with cooler

groundwater or by boiling, possibly resulting from

the abrupt pressure reduction associated with

hydrothermal brecciation.

In the "hypersaline liquid transport"

hypothesis (Fig. 9B

2

), following waning of the

magmatic vapor plume responsible for early

alteration, the lithostatic-pressured system frac

tures and the metal-bearing hypersaline liquid

ascends into the porous leached zone (Deen 1990;

White 1991; Rye 1993; Hedenquistet al.1994a).

The dominant ore-forming mechanism in this case

is mixing of the metal-bearing hypersaline liquid

with cooler groundwaters at the site of deposition,

not at depth in the meteoric water convection cell.

This hypothesis has been proposed to explain the

high salinities recorded by inclusion fluids in

several deposits e.g.,Julcani).

A part of the ore-forming components may

originate from leaching of wallrock, but both

hypotheses agree on a dominantly magmatic

source for metals, with an increase in the meteoric

water component with time. The principal

difference between the two hypotheses is in the

nature of the magmatic phase responsible for

transporting the metals into the epithermal

environment, and in the site of meteoric water

dilution. A potential contributor to ore formation

in HS deposits involves remobilization of the

metals by a meteoric-water-dominated hydro-

thermal system from a subjacent K-silicate

assemblage and porphyry-type protore, such as

that which may have formed close to the intrusion

(e.g.,Brimhall 1980). This mechanism, however,

has not been suggested as the main ore-forming

process in any of the deposits reviewed in this

study.

The three models for formation of HS ores,

assimilated here from the literature, are not

mutually exclusive; on the contrary, they may

occur in the same HS deposit as the magmatic-

hydrothermal system evolves, with complexities

arising from multiple intrusions, variations in

depth of emplacement, and changes in the local

tectonic and hydrodynamic environment. None of

the three models satisfies the overall evidence. For

example, if metals w ere supplied only by a dense,

high-salinity liquid, a relation would be expected

among estimated salinities, metal associations,

and ore grade or metal abundances of the various

deposits. Such seems not to be the case. Similarly,

if alteration and mineralization were solely the

result of interaction between groundwater and

low- and high-pressure vapor, respectively, high

salinities should not be as common as they are

unless they are explained by local boiling of dilute

to moderately saline meteoric or seawater-

dominated fluids.

SYNTHESIS

Gold, Cu, and Ag (and in a few exceptional

cases also Hg, W, Bi, Pb, and Zn) are produced

from HS deposits. As a source of Au, and because

their mode of occurrence and the potential to

overlie porphyry-type mineralization have been

widely recognized only within the past 10 to 15

years, HS deposits represent a valuable

exploration target that has been overlooked in

some regions. Most known HS deposits are young

in age, Tertiary and even Quaternary. High-

sulfidation deposits form dominantly in

subduction-related plutonic-volcanic arcs,

commonly during crustal extension. The deposits

form at a depth intermediate between the surface

and shallow (few kilometers depth) intermediate-

composition intrusions.

The intimate relationship among HS deposits,

volcanic host rocks, and oxidized magmatic fluid

derived from a degassing intrusion is supported by

the following observations: (1) the volcanic rocks

hosting HS deposits were erupted immediately

prior to mineralization, (2) the ore-forming

hydrothermal system commonly follows the same

plumbing as that of the magmatic system (i.e.,

mineralization spatially associated with domes or

volcanic conduits), (3) the isotopic composition of

hypogene sulfides (e.g., enargite and pyrite) and

sulfates e.g.,alunite and barite) commonly can be

modelled from the

34

S/

32

S of sulfur in igneous

rocks thought to be genetically related, by

equilibrium fractionation between H2S and SO4 in

solution atT-200-400 C, and (4) on the basis of

4 4 7

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oxygen and hydrogen isotopic ratios, the waters

involved in formation of HS deposits are identical

to waters in active volcanic-hydrothermal sys

tems,

in which the same HS geochemical

environment has been documented.

Ore formation in some HS deposits may

accompany acidic alteration, and recent studies of

the hydrothermal geochemistry of Au provide

preliminary evidence that this element may be

transported in HS and low-sulfidation systems as

different hydrosulfide complexes (AuHS

0

and

Au ( HS )

2

, respectively; Bening & Seward 1994;

Seward 1973). On the other hand, the presence of

moderate to high salinities in many HS deposits,

the intimate association with porphyry copper-

type deposits, and the assumptions of the most

recent genetic models (transport of Au and Cu by

either hypersaline liquid or high-pressure vapor)

indicate that chloride complexes must also be

considered for metal transport.

Most HS deposits evolve from an early period

of acidic wallrock alteration to a late period of

precious- and base-metal mineralization. Acidic

alteration is characterized by advanced argillic

assemblages and porous (leached) rock, and the

hydrothermal fluid responsible for this alteration

is dominated by high-temperature magmatic vapor

containing S0

2

, H

2

S, and HC1. Less reactive and

oxidized fluids are typically responsible for ore

mineralization. Factors such as multiple intrusions

and opening or closing of fractures (conduits)

result in variations in the temperature, pressure,

and composition of the ascending solutions.

Combined with the shallow environment of

mineralization, these conditions lead to a variety

of deposit styles (mainly replacements, breccias,

and veins) that usually occupy a limited vertical

span of 800 m at the

giant Chinkuashih deposit). The geological,

mineralogical, and geochemical evidence,

particularly the association between the orebodies

and the lateral and vertical zones of alteration,

illustrates the basic genetic condition of HS

deposits, that a magmatic fluid interacts extensive

ly with country rock and groundwaters on its

relatively short path to the earth's surface.

ACKNOWLEDGMENTS

Valuable insight on various aspects related to

this exciting ore-forming environment was gained

through discussions and field work with M. Aoki,

A. Arribas Sr., C. G. Cunningham, J. Hedenquist,

W.C. Kelly, R. O. Rye, J. J. Rytuba,and T. A.

Steven. Earlier versions of this manuscript

benefited from constructive reviews by Phil

Bethke, Andrew Campbell, Anne Thompson, John

Thompson, Peter Vikre, Noel White, and Jeff

Hedenquist, who also provided abundant

documentation on HS deposits worldwide.

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