CHARACTERISTICS OF HIGH-SULFIDATION EPITHERMAL DEPOSITS, AND THEIR RELATION TO MAGMATIC FLUID

In Magmas, Fluids, and Ore Deposits, Ed.: J.FM. Thompson, Mineralogical Association of Canada Short Course Vol. 23 (1995)

Chapter 19

CHARACTERISTICS OF HIGH-SULFIDATION EPITHERMAL DEPOSITS, AND THEIR RELATION TO MAGMATIC FLUID Antonio Arribas Jr. Mineral Resources Department, Geological Survey ofJapan, 1-1-3 Higashi, Tsukuba 305, Japan

INTRODUCTION

A consequence of the increased exploration for gold deposits during the late 1970s and early 1980s was the revision of the classification of epithermal deposits in order to account for the variations observed in styles of mineralization and inferred genetic environments. Among the numerous classifications that followed, one group of deposits clearly showed a common set of features; this deposit type is characterized by the presence of minerals diagnostic of highsulfidation states (e.g., enargite and luzonite) and acidic hydrothermal conditions (e.g., alunite, kaolinite, pyrophyllite). The terms enargite-gold (Ashley 1982), Goldfield-type (Bethke 1984, after Ransome 1909), high-sulfur (Bonham 1984, 1986), quartz-alunite Au (Berger 1986), acidsuifate (Heald et al. 1987), and alunite-kaolinite (Berger & Henley 1989) were applied to this group in reference to some of its mineralogical or inferred geochemical attributes. The term highsulfidation (HS) (Hedenquist 1987) is now widely used; the term was proposed originally to refer to a fundamental genetic aspect, the relatively oxidized state of sulfur contained in the hydrothermal system (i.e., initially S02-rich). This aspect is significant because it links HS deposits with one of the two main types of terrestrial magma-related hydrothermal systems (Henley & Ellis 1983), those associated with andesitic volcanoes whose surface manifestation includes high-temperature fumaroles and acid sulfatechloride hot springs and crater lakes. By contrast, low-sulfidation deposits form from neutral-pH, reduced (H2S-rich) hydrothermal fluids similar to those encountered in geothermal systems (Henley & Ellis 1983), with surface manifestation

including silica sinter-depositing hot springs and steam-heated acid-sulfate alteration. The main objective of this review is to summarize the characteristics of HS mineralization formed primarily within the epithermal environment, though recognizing the potential for HS conditions to occur at greater depths. Earlier studies have argued for a magmatic fluid component in HS deposits (e.g., Sillitoe 1983, 1989, 1991a; Hayba et al. 1985; Henley 1991; White 1991; Rye 1993; Hedenquist et al. 1994a), and the identification and characterization of HS deposits has contributed to a re-evaluation of the role of magmatic fluids in other types of hydrothermal systems (Hedenquist & Lowertstern 1994; Simmons this volume; de Ronde this volume). In this context, particular attention is given to the characteristics that are helpful in determining the nature of the magmatic contribution to the hydrothermal system through time and space. This review considers features of many of the deposits listed in Table 1, with locations shown in Figure 1, but is based on a selection of fourteen deposits for which the results of detailed geological and geochemical studies are available (Tables 2, and 3). For simplification, bibliographic references are not given in the text for general deposit features; these references may be fduh* in Table 1. For regional studies of HS deposits, particularly with respect to other types of magmatic-hydrothermai base- and precious-metal deposits, the reader is referred to reviews by Heald et al. (1987), Bonham (1989), Sillitoe (1989, 1991a), Berger & Bonham (1990), Camus (1990), White & Hedenquist (1990), Mitchell & Leach (1991), Mitchell (1992), and White et al. (1995).

419

A. Arribas, Jr. Table 1. Principal high-sulfldation deposits or documented prospects ordered geographically N°in 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

Deposit

References

Asia & Australasia Dobroyde, Australia Whiter al. (1995) Rhyolite Creek, Australia Raetz & Partington (1988) Temora, Australia Thompson et al. (1986) Peak Hill, Australia Cordery (1986), Harbon (1988), Masterman (1994) ML Kasi, Fiji Turner (1986) Wafi River, Papua New Guinea Leach & Erceg (1990), Erceg et al. (1991) Nena, Papua New Guinea Asami & Britten (1980), Hall et al. (1990) Motomboto, Indonesia Perelld (1994) Nalesbitan, Philippines Sillitoe et al. (1990) Lepanto, Philippines Gonzalez (1959), Garcia (1991), Arribas et al. (1995b) Chinkuashih, Taiwan Huang (1955), Hwang & Meyer (1982), Tan et al. (1993) Zijinshan, China Zhang et al. (1994) Seongsan & Ogmaesan, South Korea Yoon (1994) Nansatsu (Iwato, Akeshi & Kasuga), Japan Izawa & Cunningham (1989), Hedenquist et al. (1994a) Yoji, Japan Yui&Matsueda(1994) Teine, Japan Ito (1969) Akaiwa, Japan Akamatsu & Yui (1992), Akamatsu (1993) Mitsumori-Nukeishi, Japan Aoki & Watanabe (1995) North & Central America Northwestern Vancouver Island, Canada Panteleyev & Koyanagi (1994) Goldfield, Nevada Ransome (1907,1909), Ashley (1974), Vikre (1989) Paradise Peak, Nevada John et al. (1991), Sillitoe & Lorson (1994) Summitville, Colorado Steven & Ratte" (1960), Stoffrcgen (1987), Rye (1993) Red Mtn-Lake City, Colorado Bove et al. (1990), Rye (1993) Red Mtn-Silverton, Colorado Burbank (1941), Fisher and Leedy (1973) Mulatos, Mexico Staude(1994) Pueblo Viejo, Dominican Republic Muntean et al. (1990), Russell & Kesler (1991) South America Julcani, Peru Petersen et al. (1977), Deen (1990), Rye (1993) Castrovirreyna, Peru Vidal & Cedillo (1988) Ccarhuarso, Peru Vidal ef a/. (1989) San Juan de Lucanas, Peru Vidal & Cedillo (1988) Cerro de Pasco, Peru Graton & Bowditch (1936), Einaudi (1977) Colquijirca, Peru Vidal etal. (1984) Sucuitambo, Peru Vidal & Cedillo (1988) Laurani, Bolivia Murillo et al. (1993) Choquelimpie, Chile GiOpper et al. (1991) Guanaco, Chile Puig et al. (1988), Cuitifio et al. (1988) El Hueso, Chile Sillitoe (1991a) Esperanza, Chile Vila (1991), Moscoso et al. (1993), Cuitifio et al. (1994) La Coipa, Chile Oviedo et al. (1991), Cecioni & Dick (1992) Nevada & Sancarron, Chile Siddeley & Araneda (1990) El Indio-Tambo, Chile Siddeley & Araneda (1986), Jannas et al. (1990) La Mejicana-Nevados del Famatina, Argentina Losada-Calderon & McPhail (1994) Europe Rodalquilar, Spain SSnger-von Oepen et al. (1989), Arribas et al. (1995a) Furtei-Serrenti, Sardinia Ruggieri (1993a,b) Spahievo, Bulgaria Velinovrtc/.(1990) Chelopech, Bulgaria Bogdanov (1982,1986) Western Srednogorie region, Bulgaria Bogdanov (1982), Velinov & Kanazirski (1990) Bor, Yugoslavia Jankovic et al. (1980), Jankovic (1982) Lahoca, Hungary Baksa (1975,1986), First (1993) Enasen, Sweden HaUberg(1994)

High-sulfidation Epithermal Deposits

m

WT

C^IU-10 Western i f l P J5V-9 V *lcjk /S'fy g / Pacific

--TV? Figure 1. Worldwide distribution of high-

1I

3 - 600

!

1 I

Q.

-200

o 800

-400

I

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 of 834S suir,dMuifatt for four coexisting alunitepyrite samples (large filled circles), reference boilingpoint curves, and vertical spans of the 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 Arribas et al. 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

High-sulfidation Epithermal Deposits —.Sulfides —

— Sulfates —

T &V - 534SJS

+•

+

+

Lepanto Chinkuashlh i

Nansatsu Summltville Goldfield

—T—

Pueblo Viejo

I I

Julcani

A 3 4 SH 2 S-SO4

Temp. (°C)* H 2 S/S0 4 220 - 420

2-6

220 - 270

-

200 - 240

3

200-390

4

200 - 350

-

180-260

-

210-270

5

220-330

5

El Indto Rodalqullar

-H -10

1

1 0

1

1 1 10 &*S (%., CDT)

h20

'(mineral pairs)

30

Figure 7. Range of 834S (per mil) values for sulfides and sulfates from nine highsulfidation deposits. Also shown are the values calculated for 834S for total sulfur in the hydrothermal system (triangles), H2S/S04, and the range of temperatures determined from sulfide-sulfate mineral pairs. Solid triangles indicate deposits in which 834S£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 34 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 equilibrium, and, therefore, their overall S/ S depends on the temperature of mineralization and the 34S/32S 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 steamheated or supergene (unlikely as they contain finegrained pyrite; Jannas et al. 1990), the most likely explanation is a "magmatic-steam" (Rye et al. 1992) origin, in which the 834S 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) . Combined with the

8 S values of pyrite and enargite from the same 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; Bove et 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 534S = 2 0/

0/

± 2 'oo at Summitville, to 9 ± 2 'oo at Rodalquilar; 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 calculation done in several deposits using the S/ S values of the igneous rocks and the average 34S/32S values of sulfides and sulfates indicates that H 2 S/S0 4 in the hydrothermal fluids was generally about 4 ± 2 (Fig. 7; Rye et 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

A. Arribas, Jr.

alteration, which is characterized by a sulfate-rich alunite-pyrite assemblage. (3) Isotopic equilibrium between sulfide and sulfate in the hydrothermal solutions results, in a majority of the deposits, in reliable temperatures calculated on the basis on A SH2s-so4 (Fig- 7). Pyrite-alunite mineral pairs were used most commonly, and where sampling 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 data were available to corroborate the affinity-suggested between fluids in active volcanic-hydrothermal systems and HS deposits (e.g., Heald et al. 1987; Hedenquist 1987). Stableisotope 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 development of analytical procedures for complete stableisotope analysis of alunite, including 8 l 8 O s o 4 and 61 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

4 -> r

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 systems (e.g., Aoki 1991, 1992; Rowe 1994). Important limitations that must be taken into account for optimum use of the stableisotope 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 difference of —20 Aw between the D/H fractionation 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-known (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 zones. Oxygen and hydrogen isotopic compositions of water in HS deposits are clearly consistent with mixing between a high-temperature magmatic fluid of 8 18 0 = 9 ± l°/oo and 8D = -30 ± 20^oo and meteoric groundwaters (Fig. 8). In part because of

High-sulfidation Epithermal Deposits

0-20-

G

Alunite alteration stg.

O Q

Ore mineralization stg. Alteration/ mineralization

Subduction-related volcanic vapor Arc + crustal felsic magmas

-40-

Acidic fluids In high] sulfldation deposits .20

W W

o

Active systems (Giggenbach, 1992b)

so 6D(%.)

Volcanic Geothermal

-100

-20

i -15

1 -10

1 -5

1 0 6

18

'—r 5

10

15

20

0 (%o, SMOW)

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 8l8OSOi,); &I800H is not used because hydroxyl oxygen requilibrates with the hydrothermal fluid during cooling (Rye et al. 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 watermixing trend is displayed particularly well by the three stages of alteration/mineralization at Julcani (Deen 1990; Rye 1993): from a magmatic-waterdominated early stage of (alunite) acid-sulfate alteration (Ju; Fig. 8), through main ore-stage fluid-inclusion waters (Jut and JU2), to meteoricwater-dominated late ore-stage fluid-inclusion waters (Ju3). 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 volcanichydrothermal 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 environment similar to that of HS mineralization (Hedenquist et al. 1993). The origin of the D-enriched magmatic (endmember) 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

A. Arribas, Jr.

Hedenquist et al. 1994a; Arribas et al. 1995a). The O and D enrichment of the volcanic vapors with respect to their parent magmas (Fig. 8) is a consequence of fractionation during degassing from the melt (Taylor 1986, 1992; Matsuhisa 1992). In an alternative interpretation proposed by Deen (1990) and Rye (1993), the enrichment in deuterium is the result of reaction in a low water/rock environment between magmatic fluid of 8D ~-80°/Oo (calculated on the basis of the D/H of igneous biotite) and wallrock at magmatic temperatures down to T = -400 °C (Rye 1993), beneath the ore zone. This interpretation requires that there have been chemical and isotopic equilibrium between the magmatic fluid and country rock; however, this would likely result in neutralization of the magmatic fluid, a condition that would not favor the extreme acidity required for formation of vuggy silica at shallow depths (Giggenbach 1992a). SOURCES OF METALS

The origin of metals in HS deposits is more speculative. There is consensus among most workers (e.g., Sillitoe 1989, 1991a; Henley 1991; White 1991; Rye 1993; Hedenquist et al. 1994a) that the bulk of the ore-forming components is contributed by magmatic fluid, either directly by a magmatic vapor or hypersaline liquid that is incorporated into the hydrothermal system, or indirectly by remobilization of metals from a porphyry-type protore. The study of radiogenic isotopes so far has not provided conclusive evidence. The reason for this is a consequence of the common intimate association between the deposits and host rocks; all are indistinguishable, e.g., for Pb-isotope systematics. Lead-isotope studies at Summitville (Doe et al. 1979) and Rodalquilar (Arribas et al. 1995a) have shown that the Pb in the deposits is igneous, but these studies provide no information about the processes by which the hydrothermal system acquired the metal {i.e., leaching of the host rocks or derivation from a crystallizing magma). By contrast, in low-sulfidation gold deposits it is not uncommon to detect a component of basementrock Pb (e.g., Doe et al. 1979; Arribas & Tosdal

A1