1995 arribas
TRANSCRIPT
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A .
A rribas, Jr.
Table
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Fig. 1
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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)
Vidal Cedillo (1988)
Graton
Bowditch (1936), E inaudi (1977)
Vidal
etal.
(1984)
Vidal Cedillo (1988)
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-
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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)
High-sulfidation Epithermal Deposits
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.
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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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High-sulfidation Epithermal Deposits
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
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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
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High-sulfidation Epithermal Deposits
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
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High-sulfidation Epithermal Deposits
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)
High-sulfidation Epithermal Deposits
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
Quartz-alunitepyrite
Quartz-alunitepyrite
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,
-
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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
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High-sulfidation Epithermal Deposits
.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
435
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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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High-sulfidation Epithermal Deposits
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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High-sulfidation Epithermal Deposits
(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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High-sulfidation Epithermal Deposits
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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High-sulfidation Epithermal Deposits
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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