Chapter 20

MAGMATIC CONTRIBUTIONS TOLOW.SULFIDATION EPITHERMAL DEPOSITS

Stuart F. SimmonsGeothermal Institute, The University of Auckland,

Private Bag 92019, Auckland, I,{ew Zealand

INrRonucrron

Low-sulfidation epithermal deposits form at<300'C and depths <2 km within meteoric-water-dominated hydrothermal systems that aregenetically linked to magmatism at convergentplate boundaries. Despite this strong associationwith intrusions causing convection, the occurrenceof magmatic fluids in the low-sulfidationepithermal environment is difficult to detectbecause of scant diagnostic evidence. There arethree reasons for this:

1 . M a g m a t i c s i g n a t u r e s m a y b e m a s k e d b ywater-rock interaction along the flowpathseparating a magma from a suprajacent low-sulfi dation environment.

2. Magmatic contributions likely wax and wanewith time, occurring in some cases in sharppulses. Consequently, their preservation maybe concealed within complex inter-mineralbanding and/or intra-mineral zonation, atcentimeter to micrometer scale.

3. Waning hydrothermal activity may dilute orerase magmatic signatures.

Despite these geologic processes, the resultsof a few detailed studies of low-sulfidationepithermal deposits indicate that magmaticcontributions do exist and are possibly importantto ore formation. The salinity, and gas andisotope compositions of fluids in activegeothermal systems of low-sulfidation charactersubstant iate this point of v iew.

In this paper, I review the evidence formagmatic contributions in the formation of sixworld-class Iow-sulfidation epithermal deposits.drawing on the fluid-inclusion and isotopic studies

published over the last twenty years. Along withbeing well-studied, the geologic settings andmetal contents of the deposits are diverse,spanning the spectrum of low-sulfidation deposittypes. The focus is on big deposits because,assuming that a cause and effect relationshipexists, the potential for finding evidence ofmagmatic contributions and their possible link toore genesis is greatest here. Accordingly, thewell-studied but small epithermal deposits, suchas Creede. USA. are not discussed. Theframework for interpreting the data from thesedeposits, however, is based on studies of activegeothermal systems, where the heirarchy ofphysical and chemical processes that affectthermal fluids can be directly assessed in amodern geologic context. Therefore, I first reviewthe evidence of magmatic contributions in activegeothermal systems, referring specifically to datafrom the Taupo Volcanic Zone. In discussinglow-sulfidation systems, I will emphasize theimportance of time scales in which both long- andshort-term magmatic contributions are indicatedand will discuss their relevance to ore-formation.

MAGMATIC Con poNENTS IN ACTIVEGnorrrnRIrAL SYSTEMS

Epithermal deposits form in the shallow partsof hydrothermal systems where temperature,pressure, and chemical gradients favor efficientprecipi tat ion of metals through boi l ing or mixing,from a deeply derived fluid (e.g, Heald et al.1987; White & Hedenquist 1990). Studies ofactive geothermal systems (e.g., Giggenbach1981, 1987, 1988; Hen ley & E l l i s 1983;Hedenquist 1986) indicate that the composition ofthis deep fluid is shaped by its circulation history:

455

S.F. Simmons

l) incorporat ion of magmatic volat i les into largeinputs of meteoric waters at deep levels; 2)water-rock interaction at intermediate levels; and3) boi l ing (stearn and gas loss) and mixing(typical ly di lut ion) at shal low levels (Fig. l ) . Theaqueous and gaseous components that precipitateas minerals within epithermal deposits, or that aretrapped as inclusions in minerals, can therefore beinf luenced by one, or a combinat ion of threernajor sets of processes. By considering only thosecomponents which are non-react ive orconservative to the gas and liquid phases, themasking effects of shallow and intermediateprocesses can be deterrnined.

Table I lists the components from whichmagmatic signatures in epithermal environmentscan be interpreted, based on studies of volcanicgases, geothermal fluids, and hydrothermalminerals. These signatures involve a rat io ofcomponents (N2-Ar-He) or isotopes (hel ium,oxygen, hydrogen). Only for chlor ide (or sal ini ty)can a magmatic signature be inferred from itsconcentrat ion alone. The most diagnost icsignatures are indicated by helium isotopes andN2-Ar-He, both of which occur in smal l to traceamounts, though analysis of these components inepithermal minerals has not been widely appl ied.In contrast, interpretation of the more abundantand commonly determined chloride (reported aseq. wt. % NaCl from ice melting temperatures off lu id inclusions), and oxygen and hydrogenisotopes, provide permissive evidence throughconstruction of internally consistent argumentsthat establ ish basel ine andlor end-member com-positions to interpret data trends. These conceptsare elaborated below by considering thesecomponents and their magmatic signatures ingeothermal fluids from the Taupo Volcanic Zone.

Geothermul Systems in the Taupo Volcanic Zone

(TVZ)The

'lYZ contains about twenty equispaced

geothermal systems associated with calc-alkalinemagmatism (Fig. 2). About half of these systemshave been dr i l led to depths of 500 to 2800 m; thedeep fluids discharged from these wells have beenanalyzed for rnajor aqueous and gaseous

components, plus hel ium-isotopes, 6180 and 6D

4 5 6

Figure l. Schematic diagram showing theenvironment for low-sulfidation epithermal minerali-zation within a magma-related hydrothermal systemwith low topographic relief (after Henley & Ellis1983; Hedenquist 1986).

(Table 2; Hedenquist 1986; Giggenbach 1995).These data were corrected for effects due to steamloss and dilution and therefore represent thecomposit ion of the deep I iquid pr ior to boi l ing.Host rocks comprise a sequence of Quaternaryvolcanics of predominantly rhyolitic compositionthat unconformably overlie weakly rnetamor-phosed graywacke basement of Mesozoic age(Cole 1990); the relat ively uniform composit ionof the stratigraphy here helps in assessing thecontributions of rock-leachable constiuents togeothermal waters. Direct evidence of intrusionsis rare (Browne et al. 1992), but their existence isinferred from the high heat flow associated withgeothermal activity, and the distribution andoccurrence of volcanic events. Fluid samples fromactive andesitic volcanoes at White Island andNgaruhoe, at the north and south ends of the TVZ,respectively (Fig. 2), provide constraints on theend-member composit ion of magmatic gases inthe region.

The TYZ geothermal systems also containgeochemical environments that are analogous tothose deduced to be responsible for forming low-suffidation epithermal deposits (e.g., Henley1985). This perspective is substantiated byoccurrence and zonation Datterns of hvdrothermal

C\-N-

III

E

@.c

A ,. t ,

, i : , t ,, : . :: nagru :

ore fotming environment

L ow- s u I /idati on Ep i t herma I Dep os i ts

Table L Components indicating magmatic signatures in low-sulfidation epithermal depositsmagmat ic

components s ignature'r

r"/H" I l i Ra> I

N, Ar -Hc I l c iAr . . i

N . , / l l c . . "10

6'to 8 to l0 "/o.,

C l l conccn- >10.000 mg/kgtration

evidencediagnostic

diagnostic ( ' l ) -

chcck fbr manlle-hcl ium isotopesignature

pcrmissive - check

with 61,)

6l) -20 to -80 "/". , pcrmissive for datatrends assesscd in

conccrt rvi th 6r8oval ucs

so u rcefluid inclusions

f lu id inc lus ions

quarlz. calcite (requiresestimatc of cqui lbrat iontemperaturc, e 9., 7-r):f luid inclusions

fluid inclusions

temporal resolut ionpoor to moderatc - rcquires bulk anal l 's isof rnaterial potcntial l l ' containing rnutiplcgenerations of f luid inclusions

good to moderate - requires anal l sis o1'matcrial potential ly conlaining mutiplegenerations o1'f ' luid inclLrsions: individualinclusions sampled bv thermaldecrepitat ion

good-rcquires a lew tens of mil l igrarnsmatcrial

poor 1o rnodcratc - requircs bulk lnalvsisof material po{cntial iy containing rnutiplcgcncr r t ions r ' ! - l l u id inc lus ions

excel lent - individual 1' luid inclusions canbe measured

permissivc only 1br f luid inclusionsrclat ivcly young crustlacking sonnatcbrint-s or cr aDoritcs

alteration. the state of hydrotherrnal fluid-mineralequi l ibr ia, occLlrrence of "ore-grade" concen-trat ions of precious and base metals inprecipi tates, and relat ively high precious-metalf luxes (e.g., Brown 1986; Brow'ne 1969; Browne& El l is 1910 Browne & Lovering 19'73;Hedenquist & Browne 1989; Hedenquist &.Henley l985at Henley 1985; Krupp & Seward1987: Simmons et al . 1992; Simmons &.Christenson 1994; Weissberg 1969). The compo-sitions of deep geothermal waters are dominatedby concentrat ions of chlor ide (up to 2500 mg/kg)and dissolved carbon dioxide (up to 20,000mg/kg), occurrence of reduced aqueous sulfurspecies ( t{2S, HS-) and near-neutral pH (e.g., El l is1919), with their equi l ibr ium rnineral assemblagecomprising quartz, orthoclase ("adularia"), il l ite.chlor i te, pyr i te and calci te. + epidote andwairaki te (Giggenbach 1988). These deeplyderived waters also transport relatively largequantities of precious and base metals, whichprecipitate in hot springs, fractures, and vugs, andin wel lhead piping. Of these, the spectacularsulf lde scales, containing up to 6 wt. 0% Au and 30

wt. o Ag, on wellhead back-pressure plates atKawerau, Broadlands-Ohaaki, and Rotokawademonstrate: 1) that high concentrations ofaqueous Au and Ag (-l to 2 pg Au/kg; 8 pgAg/kg) in chloride waters are possible due tobisulfide complexing (Seward 1973 ). and 2) thatboiling from about 260 to 180 'C is an efficientmechanism to precipitate Au and Ag (Brown1 e86) .

Metal fluxes in these three eastern systems(i.e., Kawerau, Broadlands-Ohaaki, andRotokawa; see Fig. 2) appear higher than those inthe central and western systems due to the highergas (CO2 and H2S) concentrations (Henley &Hedenquist 1986). The mass flow for the pre-exploitation state of Broadlands-Ohaaki isestimated at 100 kg/sec; thus, about 6500 yearsare required to f lux I mi l l ion oz Au (0.03 x 10"kg) thro"ugh the upflow zone, which covers about10 km' (Brown 1986) . Shor te r per iods fo requivalent Au flux are inferred for both Kawerauand Rotokawa because both aqueous sulfideconcentrations and mass flows are higher (Henley& Hedenquist 1986; Krupp & Seward 1987).

S.F. Simmons

White lsland 6

of Plenty

^1(;Tarawera

,*L ffin.""n"M a n g a k i

- {A"Waiman^"fwuiotuo,

lod c""ooro"

Ngatamar ik iB R O A D L A N D S /O H A A K I

W a i r a k e i j

GEOTHERMAL AREAS

Delineated bY dril l ing

Del ineated bY geophYsics

Other thermal areas

VOLCANOES

Geothermal systems remain active for several tens

of thousands to several hundreds of thousands of

years (e.g., Browne 1979; Henley & El l is 1983),

indicating that periods of metal deposition at

epithermal depths can be relatively short-lived and

still account for very large, high-grade orebodies

(Henley 1985), based on the concentration of gold

and silver directly measured in the high-gasTYZ

systems.

Helium IsotopesThe large differences in helium-isotope

compositions for atmospheric, upper-mantle, and

crustal or radiogenic sources, makes He a useful

tracer in volcanic and geothermal gases (Fig. 3).

The atmospheric 3He/He rat io (Ra) equals 1.39 +

0.01 (Clarke et al. 1969; Mamyrin et al. 1969) and

is used as a reference standard to which all sample

4 5 8

Figure 2. Location map ofTVZ geothermal systemsdescribed in the text (from

Hedenquist 1990).

isotopic ratios are compared in the form R/Ra.

The MORB R/Ra value of 8-9 is determined from

gas analyses of mid-ocean ridge basalts and axial

vent hot spr ings (Craig & Lupton 1981). Deep

undepleted mantle may have much higher R/Ra of

15-30 (Lupton 1983). The continental crust R/Ra

value of <0.1 results from the decay of uranium

and thorium (Gerling et al. 19'11; Morrison & Pine

1es5) .Volcanic and geotherrnal gases from the

circum-Pacific have helium-isotope compositions

that range from I to 8 times Ra, indicating that

they contain a significant component of mantle-

derived helium (e.g., Giggenbach eI al. 1994;'

Hilton et al. 1993; Sano et al. 1985, 1981;

Torgersen & Jenkins 1982; Torgersen el al. 1982).

Mantle helium is thought to rise through the crust

with buoyant melts, exsolving with other gases at

AtmosphereRlta=1

Figure 3. Helium isotopic compositions from themain terrestrial sources, including atmosphere, uppermantle. and continental crust.

relatively shallow depths (<10 km) where magmascrystallize in subvolcanic chambers. Solid rockscontain insufficient 'H. to account for its rela-tively high concentrations in geothermal fluidsdue to leaching. Therefore, mantle helium mustindicate a direct magmatic input and is consideredhere to be the single most diagnostic indicator ofmagmatic contributions to geothermal fluids andlow-sulfi dation environments.

In the TVZ, geothermal gas samples rangefrom 3.5 to 8 R/Ra and vary independently ofCOz, Cl, or other component concentrations(Giggenbach 1989; see Table 2). This l ikelyresults because He is a noble gas, and because theprocesses affecting He-isotope ratios aredecoupled from those affecting other components(Giggenbach et al. 1994; Sano et al. 1981;'Torgersen et al. 1982).

Table 2. Chemical characterist ics of TVZ geothermalf luids (Hedenquist 1986; Giggenbach 1995)

-4.5 -40

-3.75 -3 I- a . J - + J

-6 -43

-6 5 -45

-4 75 -38

-2.5 -40

-3 s -10

Low-suUidation Ep ithermctl Depos its

Nit r o g e n-He I i u m-Arg o nThe relative concentrations of N. Ar. and He

sampled from active volcanoes associated withmantle hot-spots, crustal rifts, and convergentplate boundaries plot in two distinct groups (Fig.4) that define mixing trends involving basaltic,andesitic, and meteoric water end-members(Giggenbach 1992a). The "andesitic" end-memberis characterized by N2/He values that range from1700 to 5000. By contrast, the "basaltic" end-member is characterized N2/He values that rangefrom l0 to 220, though the relative enrichment inhelium may alternatively suggest a crustal source(Giggenbach 1986). Values of He/Ar for thesetwo source regions are close to the mantle ratio ofabout 3. Air-saturated meteoric groundwater(N2/Ar value about 50) forms the common end-member to both trends and indicates the influenceof the atmosphere (N2/Ar 82) on gascompositions. Distinct enrichments in CO2/Hevalues for volcanic gases also exist for "andesitic"

versus "basaltic" gases, implying that both N2 andCO2 are derived from the same source, whichGiggenbach (1992a) speculates is subductedsediment.

The N2-Ar-He signature of geotlrermal fluidsfrom high-temperature systems in Japan (Chiba19911' Ueda et al . 1991; Yoshida l99l) , NewZealand (Giggenbach 1992a), and the Philippines(Giggenbach & Poreda 1993) overlaps withvolcanic gas data and indicates the existence ofdeep inputs from either basaltic or andesiticsource regions in these systems. As nitrogen isrnostly non-reactive in geothermal fluids, as areargon and helium, the relative concentrations ofthese three gases complement helium-isotoperatios as a potential tracer of deep magmaticcontributions. For example, geothermal systemslocated along the eastern margin of the TYZltaveN2-Ar-He compositions that plot between theandesitic and meteoric end-members, whereaswestern TVZ geothermal systems have N2-Ar-Hecompositions that plot between a basaltic andmeteoric end-member (Fig. 4). These resultsindicate the incorporation of magmatic gases intoTVZ geothermal fluids. In addition, thesesignatures correlate with high and low total-gasconcentrations for the "andesitic" and "basaltic"

systems, respectively, and indicate that the

l . Broadlands 970Ohaaki

2. Kawerau 8l 0

3. Rotokawa 850

4. Wairakei 1700

5. Mokai 2400

6. Waimangu 5t l5

7. Waiotapu 1250

8. I 'okaanu 2500Wa ih i

9. Whitc Is land >10,000

4s50 6 . l

3 1 6 0

43 80

I 820

t27 5

l 8 7 5

t620

2000

1720

7 . 1

u.27 . 0

7 . 5

7 .0

4 .0

6 .6 + 7 . l

Continentalcrust

R/Fa<0.1

1 1

4s9

S.F. Simmons

Geothermal

Volcanic

2OO Ar

And BasO D

50ASWr.'- 20

Figure 4. Relative concentrations ofN, Ar, and He in gases fiomandesitic (And) and basaltic (Bas)volcanoes around the world andgeothermal gases fiom the

'fVZ, all

having hel ium isotopic rat ios ofR/Ra >4 (Giggenbach 1992a).

lOOO He

magmatic gas components, when present, result in

a much higher gas concentration than that of

meteoric or crustal source (Giggenbach 19924'

Giggenbach et al 1994).

Oxygen and Hydrogen IsotoPesOxygen and hydrogen are the principal

components of geotherrnal waters, and their

isotopic compositions plot close to that of the

local meteoric composit ion (Figs. 5a,b), indicat ing

this is the primary source of water (Craig 1963).

The commonly observed positive enrichments in

6180 and 6D result f rom one, or a combinat ion, of

the three main processes affecting the composition

of geothermal f lu ids (Fig. I ) .At shal low levels, the isotope composit ions of

the deep fluid can increase upon ascent to the

surf'ace through steam loss associated with

boi l ing. Both cont inuous and single-step modes

of steam loss have been quantitatively assessed,

assuming adiabatic cooling, and the results

indicate that single-step separation produces the

maximum isotope enrichments (Truesdell e/ a/.

1977; Giggenbach & Stewart 1982). For example,

single-step steam loss for a l iquid which boi ls

from 260 to 100 "C produces positive shifts of

about 1.5 "/oo

6180 and 10 "/oo

6D (Fig. 5c).

Subtracting these eff-ects from the compositions of

surface waters sarnpled frorir wells and springs

indicates that deeper level processes tnust beconsidered to explain tlre isotope compositions ofthe pre-boiled parent waters.

Intermediate-level processes involve isotopeexchange through water-rock interaction aI

elevated temperatures that depletes the 6l80 and

6D compositions of fresh rocks and enrich the

6180 and 6D compositions of geothermal waters(Craig 1963; Taylor 19731' Blattner 1993). Asfresh rocks contain large amounts of oxygen (up

to 46 wt. o/o) and relatively small amounts of

hydrogen (< 0.2 wr.o/o), the 6r80 composition ofeither the rock or meteoric water is rnore easilyshifted than the 6D composition as controlled bythe bulk water-rock ratio and temperature-dependent mineral-water fractionation factors.These isotopic shifts have been modelledquantitatively, assuming sirnple closed and open

systems (e.g., Ohmoto & Rye 1914, Field &Fifarek 1985; Criss &

-faylor 1986), and more

complex dyrramically evolving infiltratiorr fronts(e.g., Blattner & Lassey 1989; Blattner 1993).

Giggenbach (1993) has also constructed a simpleisotope-exchange model based on hydrothermalalteration as represented by the followingreaction:

Ca-f 'eldspar + K-f 'eldspar + l l ,O + COr -+ K-mica + 2 quartz-r calcire (1)

Zsn\ \/ x w a \ \

/"c. ?-. \{Qi13j"'ir

-7- a

6 D % "

- 1 5

6 D % "

0

-20

-40

-ou

-80

Giggenbach pointed out that pure water on itsown is relatively inert and unlikely to breakstructural bonds and drive the hydrothermalreactions required for isotopic exchange, hencethe relevant inclusion of CO2 as the main acidspecies, promoting the hydrolysis of feldspars(see C iggenbach 1981, 1984, 1988) . Thus , one

Low-sulfidation Epithermal Depos its

6t"o %'u

mole of CO2 is consumed in converting Ca- andK-feldspars to K-mica. Accordingly, minimumwater-rock ratios involved with isotopic exchangewil l range between 0.15 and l , given thatH2O1CO2 rnole ratios in volcanic fluids are > 5and in geothermal fluids >30. Using these valuesin calculations (see Giggenbach 1993) indicates

-2- 4-6

6180 %,o

Figure 5. The oxygen and hydrogen isotopic compositions of geothermal and volcanicfluids: (A) geothermal systems associated with andesitic volcanism (from Giggenbach1992b); (B) TVZ geothermal systems in Table 2; (C) the oxygen and hydrogen isotopicenrichments in l iquid due to adiabatic single steam loss from 260 to 100 oC (Giggenbach &Stewart 1982). The water-rock interaction curve in (B) assumes equil ibrium isotopic

exchange between local water (6'tO : -7 o loo and 6D : -45 o/oo; and K-mica at 300 'C (see

Giggenbach 1993); dots represent water-rock weight ratios.

O local groundwaters

O geothermal dischargesI volcanic condensates

. SMOW

Whitelsland

.oG \->

rr'i'x\$g2.'P

1. Broadlands Ohaaki 5. Mokai

2. Kawerau

3. Rotokawa

4. Wairakei

6. Waimangu

7. Waiotapu

8. Tokaanu-Waihi

aoitino pc

,*€ c

S.F. Simmons

max imum pos i t i ve enr ichments o f about 9 o loo

6'nO and 6.5 u/nn

6D due to water-rock interaction(F ig . sb) .

Water-rock interaction has long been thoughtadequate to explain the isotopic enrichmentsobserved in geothermal fluids. However, recentscrutiny of the isotopic compositions of geo-thermal waters associated with arc volcanism inthe circum-Pacific indicates a common input ofmagmatic water to geothermal systems (Fig. 5a).For these geothermal waters, both 6180 and 6Dvalues are higher in comparison with those inlocal meteoric water. The straight lines connectingIocal meteoric and geothermal waters increase inslope with increase in latitude. The lines formtrajectories that point to a common end-membercomposition of about 9

"/oo 6'tO and -20 o/oo

6D,close to the composition of high-temperaturedischarges from nearby andesitic volcanoes (Fig.5a). This so called "andesitic" water (Giggenbach1992b) has a 6D value about 45 + l 5

o/on heavier

than that which was previously estimated formagmatic waters on the basis of analyses ofamphiboles arrd micas from eroded plutons(Sheppard et al. 1969 Taylor 1986). The lighter6D signature results from isotope fractionationbetween water in a melt and water vapor exsolvedduring crystal l izat ion (Taylor 1986), and theactual composition of water in an undegassed meltis l ikely intermediate in composit ion (Taylor1992). The trends in Figure 5a thus indicate thatrnagmatic water contributes to some geothermalwaters, in proportions up to 20 %, since isotopeenrichments due to water-rock interaction alonecannot explain the change in slopes of meteoric-geothermal water trajectories as a function oflatitude (Giggenbach 1992b, I 993).'fhe

stable-isotope compositions of parent watersin TYZ geothermal systems (Fig. 5b) i l lustratethat resolution of the relative contributions ofintermediate (water-rock interaction) and deep-level (magmatic and meteoric water mixing)effects is difficult for many cases because thestrengths of their respective signals significantlyoverlap. Hence, stable- isotope composit ionsprovide permissive, but not diagnost ic, evidenceof a rnagmatic contribution to deep geothermalwaters.

ChlorideChloride concentrations for geothermal waters

in systems associated with arc magmatism,isolated frorn seawater recharge and lackingevaporites at depth or a source of connate brines,typically range from about 500-2500 ppm forrhyolite-related systems (e.g., Table 2) and 6000to 12,000 ppm for andesite-related systems(Hedenquist & Henley 1985b). As is the case forstable isotopes, the chloride concentrations ofgeothermal waters can be related to one, or acombination, of the three main processes affectingthe composition of geothermal fluids (Fig. I ).

Because the meteoric recharge for deepconvective systems stafts as almost pure water,the concentrations of chloride measured ingeothermal waters must be derived throughprocesses other than boiling. The magnitude ofboiling effects, however, can be calculated todetermine the parent composition (see Henley eral. 1984). For example, a rising liquid that boilsadiabatically from 300 to 100 oC produces about40 Yo steam by weight, and concentrates chloridein the l iquid by a factor of about 1.7; in this case,1700 mg/kg chloride in a water discharged at thesurface indicates its deep pre-boiled parentcontains 1000 mg/kg chlor ide. Such enrichmentsare l ikely to be close to the maximum, given thatthe temperature of first boiling in most geothermalsystems rarely exceeds atlout 300 'C.

Experimental results indicate up to severalhundred mg/kg aqueous chloride can be derivedthrough leaching of crustal rocks by pure hotwater (El l is & Mahon 1964. 1967. Fig. 6).Accepting a minimum bulk water - rock weightratio of 0.1 for geothermal systems, maximumchloride concentrations of up to several thousandmg/kg are possible and overlap with the range ofcompositions observed in active geothermalsystems.

Alternatively, magmatic contributions canaccount for chloride in geothermal fluids, as isindicated by the concentrations of up to a fewweight percent in high-temperature volcanicemissions (e.9., Hedenquist & Lowenstern 1994)and high-sal ini ty f lu id inclusions (>35 wt. %NaCl) associated with igneous intrusions (e.g.,Roedder 1984). The processes relating to tlre

462

ct)

CDtr: 200o

0.075Eio

O.O5O o.:g

Eo.o,, Hz

Figure 6. Concentrations of aqueous chlorideexperimentally teached from different rock types for aperiod of 14 days at >500 oC, using pure water andwater-rock ratios * I (Ellis & Mahon 1964, 1967).

transfer of magmatic chloride into the overlyinggeothermal convective cell are not well-

understood but are likely to be more complex than

can be assumed from simple mixing (e.9.,

Fournier 1987), and this makes magmatic inputs

difficult to quantify.Deep pre-boiled TVZ geothermal waters

contain up to about 2500 mg/kg chloride, whereas

the magmatic water from White Islands contains>10,000 mg/kg (Table 2). Even so, the source of

chloride in these geothermal waters, whether it be

derived from water-rock interaction or direct

magmatic input, cannot be distinguished on the

basis of the chloride concentration alone. The

range of chloride concentrations in parent

geothermal waters feeding systems across the

TYZ (Table 2), which varies by a factor of 5

despite the passage of the waters through similar

country rocks, suggests a source other than simple

rock leaching. The variation in B/Cl ratios

(Giggenbach 1995) across theTYZ, and chlorine-

isotope studies (Hedenquist et al. 1990), instead

point to a deep igneous origin for chloride. In the

absence of these parameters, only from extreme

enrichments in chloride (e.g., >10,000 mg/kg) can

a magmatic source be inferred.

MAGMATIC CON PONNNTS IN EPITHERMALORn Dnpostts

Low- sulfidat ion En it her mal Dep os its

ter ist ics of s ix wel l -studied, world-class, low-sulfidation epithermal deposits for which evidenceof magmatic contributions exists. These depositshave been mined for precious and base-metals(Fig. 7), which occur as open-space f i l l ings inveins and stockwork structures or within porespaces among sedimentary clasts. The geologicsequences associated with these deposits arediverse, but the rocks that host orebodies are madeup of volcanics (Antamok-Acupan, Comstock,Emperor, Tayoltita) and/or underlying and olderbasement units containing sedimentary andmetasedimentary lithologies (Fresnillo, Hishikari).None of the deposits have rock types in theirstratigraphy that are likely to contain evaporites orconnate brines.

The main geologic feature common to all ofthese deposits is their close spatial and temporalrelationship to regional magmatism associatedwith convergent movements along a plate margin.The composition of nearby igneous rocks may ittfact have a genetic influence on the metal contentsof ores (Sillitoe this volume), as is postulated fbrEmperor, where gold-silver-telluride ores arerelated to alkalic mafic igneous compositions(Richards this volume). lgneous intrusions occurat all of the deposits except Hishikari, whichformed at shallowest depth. At Antamok-Acupan,

A=Antamok-AcupanC=ComslockE=EmperorF=FresnilloH=HishikariN=NZTvZT=Tayoltita

Cu+Pb+Zn Ag (x100)

Figure 7. Temary plot of relative metal contents of

ores produced from low-sulfidation deposits in Table

3. The metal ratios of TVZ geothermal precipitates

are similar to the metal ratios of ores from Comstockand Tayoltita.

400

100

ox(!

3o

-9 E 9 9 1 o = =E E E E E g Hc - o r : ! c !

Au (x100)

Table 3 the salient geologic charac-

463

S.F. Simmons

Table 3. Geologic characterist ics of low-sulf idation epithermal depositsAssociatedigneous rockDeposi t Au

Antamok- 0.5Acupan,Phi l ippines

Comstock 0.28 6.56l .ode, USA

ApproxPrincipal arealmetals extent

Countryrocks

volc-intru-diatreme-sed intermediate(meta)

Rct'ercnces

omsti e/ al. 1990Cookc & Bloom1990; Mi tchel l &l , each l 99 l ;Sawkins et al. 1979Vikrc e/ a/ . 1988;Vikre 1989

Ahmad et al. 1987;Andorson & Baton1990; Eaton &Serter f ie ld 1993;Kwak 1990,Setterficld et al1992Gsmmelf et a l .1988.I .anget a l . 1988;Macdonald et a/.I 986; Ruvalcaba-Ruiz & Thompson1988 ; S immonsI 991lzawa et al. 1990

Smith el al. 1982

Orebodies AgeMa

Empcror,I'rj i

0 . I I 0 . 0 4

Ag-Au 9 km2

Au-Ag- 6 km2Te

volc-intru-metased-metavolcvolc- intr(sed)

sed-volc-intru

intermediate-l'elsic

mal ic (a lkal ic)

intermediatc-fe ls ic

5 (?) Au-Ag veins, stockdisseminations;pre-exis ing high-S & porph min

velns; pre-existing high-S &porph minveins, stockwork;pre-existinghigh-S & porphmin

veins, stockworkmantos, chimneys

F resn i l l o , 0 . 015 10 .0Mexico

I l ishikar i , 0 25 0. l 5Japan

Tayoltita,Mexico

Ag-Pb- l0 km2Zn

Au-Ag I km2

15.0 Ag-Au 4 km2

-l scd-volc intermcdiale

-29

0 3

vetns

veins volc- intru intermediate

rAbbreviations: volc:volcanic, high-S & porph min=high-sulfidation and porphyry-style mineralization, intru:intrusive, sed:scdimentary,

meta:metamorphic.

Comstock, and Emperor, prolonged periods ofhydrothermal activity produced spatially relatedporphyry and high-sulfidation epithermal styles ofmineralization that preceded the emplacement oflow-sulfidation ores (Comsti et al. 19901. Cooke &Bloom 1990; Mitchell & Leach 1991; Setterfieldet al. 1992; Sillitoe & Gappe 1984; Yikre et al.1988: Vikre 1989). Their occurrence, however,does not influence the discussion. below. on low-sulfidation ores.

Characteristics of the fluids associated withmineralization are listed in Table 4. Finecentimeter-scale mineral parageneses are the basisfor fluid-inclusion and stable-isotope studies,except for Comstock, for which the samplesinvestigated were obtained from universities,museums, and private collections (Vikre 1989).Hydrogen-isotope analyses are available forAntamok-Acupan, Comstock, Emperor, andFresnillo, but helium-isotopes and N2-He-Ar gas

analyses are limited mostly to Fresnillo. Fluid-

464

inclusion data and hydrothermal minerals, withthe occurrence of sulfides and "adularia"-il l ite inevery deposit, indicate conditions of minerali-zation of 300 to 150 oC from fluids of low-sulfidation character, having reduced aqueoussulfur species and near-neutral pH, similar to thedeep fluid compositions of TVZ geothermalsystems. These similarities provide the basis forinterpretat ions below.

ChlorideOf all the components that could be used to

indicate a magmatic source, chloride, or salinity,is the easiest to measure by observing ice meltingtemperatures in fluid inclusions. Given carefulfluid-inclusion petrography combined with thesemeasurements, salinity variations provide theclearest sign of change in fluid composition withtime. The biggest drawback is the uncertaintyregarding the salinity, as the ice meltingtemperature (7.) between 0 to - 1.5 'C may also be

Low-sulfi dat io n Ep i I he r ma I D e p os i I s

Table 4. Characteristics of mineralizing fluids in low-sulfidation epithermal depositsSalinity l{ydrothcrmal

'Hei He

Deposit 'I '"Cl wto/o mineralsr 6r8oqtz"/oo 618c)"rooloo 6Dn,o

o/oo (R/Ra) Rcf'crcnces

82 et a l . 19 '79. Simmons 1986n cupan,Phil ippincs

Comstock 250-300 l-7Lodc. [JSA

Bmperor. 160-300I U l

i l -chl -ep-mag

qtz-calc-sul l ' -ad -1. I to 8.9 -7.8 to L9

i l -chl -mont

4.5-7 qtz-calc-sul f ' -ad 16.9 to 6.5 to 12.9

i l -chl - ro-mont 20.5

Frcsni l lo , 200-300 l - l 5

Mex i co

l 3 . 9 t o 3 . 9 t o ' 7 . 7t 7 . 4

-69 to -r 3 3

-10 to -5 8

- 3 0 t o - 1 t o 274

O'Nei l & Si lberman 1974;1 'aylor 1973; Vikre et a/ . l98t ' i ;Vikre 1989

Ahmad et al. 1981 , Anderson &Eaton 1990; Eaton & Setterlield1993; Kwak 1990; Set ter l le ld etal 1992

Macdonald et al. 1986,Ruvalcaba-Ruiz &' l 'hompson1988; Simmons et a l . 1988.Simmons, 1991

lzawaet al. 1990 Matsuhisa &Aoki 1994

Churchi l l 1980; Clarke &' l ' i t ley1988; Conrad et al. 1992 Smith

qtz-calc-sul f'-ad-i l -chl -mont

I I i sh i ka r i , I 50 -250

Japan

l ayoltita, 250-280Mexico

derived from dissolved gases, mainly COz

(Hedenquist & Henley l9S5b). Thus, the only data

for which saline solutions can be confidently

interpreted are those indicating >2.5 eq. wt. o

NaCl. Furthermore, because of the vagaries

associated with fluid-inclusion trapping. ice

melting studies require a minimum of five to ten

measurements from a few spatially distributed but

paragenetically constrained samples to be certain

of the range of fluid comPositions.The salinity data for the six deposits are

plotted in Figure 8 and show a range from <1 to

15 eq. wt. o NaCl. The data for Hishikari and

Tayoltita indicate that the mineralizing fluids were

dilute waters, which is typical of the majority of

gold-silver epithermal deposits (Hedenquist &

Henley 1985b); note that for Tayoltita a few fluid

inclusions containing >4 wt. % NaCl exist , and

Conrad et al. (1992) suggest that these represent

inputs of magmatic fluids. In contrast,

m ineral iz ing f lu ids for Emperor, Fresni l lo,

Comstock, and Antamok-Acupan were saline(>2.5 eq. wt. yo NaCl), and the chloride in these

systems may have a magmatic origin. Only for

Fresnillo. however, can saline fluids be

<l qtz-calc-sul f ' -ad 7.3 to 9.8 -5.7 to -0.1

i l -chl -mont

0-13 qtz-calc-sul l ' -ad- 3 l to 8.0 -5.8 to I l

abil-chl-aot-eD et al 1982

adu la r i a . ca l c_ca | c i t e . ch |_ch |o r i t e . ep : cp ido te . i | _ i | | i t e .

mag=magnetitc. mont- montmorillonite, q1z:quartz, ro-roscoelite, sulFsulfides, zeol=zeolite

genetically linked to mineralization, with Ag-Pb-Zn transport being favored by chloride complexes.

Figure 9 shows some of the results of a fluid-inclusion study associated with the infil l ing of the-1 m wide Santo Nino vein in the FresnilloDistrict. Spatial and temporal variations in

compositions indicate episodic injections of brines(up to 12 eq. vrt. % NaCl) into fractures otherwisefifled with relatively dilute fluids (-3 eq. wt. yo

NaCl). Brines are closely associated with sulfidemineralization, occurring primarily in sphalerite-hosted fluid-inclusions, whereas the low-salinityfluid-inclusions are hosted by barren quartz and

calcite. This relationship is found elsewhere in the

Fresnillo district, indicating that mineralizationwas coincident with the repeated introductiorr of

saline liquids. The high salinity is interpreted to

have been derived from a brine reservoir of

magmatic origin, which was situated beneath a

dilute geothermal convection cell (Simmons

1991). The rocks in the distr ict most ly consist of

a highly deformed graywacke-argill ite sequence

overlain by younger rhyolite volcanics, unlikelyhosts for connate brines or evaporites. Nor can

simple rock leaching account for the slrarp

465

II

II

S.F. Simmons

ctl:E

40,000 E()

cr 1n

{B6 o(uz

E p E b t gj i = a 1 l i o =

= g 3 F E bl g i l d E E Er F j u r n ' ! -

g

Figure 8. Range of salinit ies determined from fluid-inclusion ice melting measurements (7.) for epithermaldeposits (Tables 3, 4). The shaded region shows the 1,,range for which the freezing-point depression due todissolved salts or carbon dioxide cannot bedist inguished (Hedenquist & Henley 1985b).

Stage I

variations and high salinities over time. Furtherevidence that the high salinities at Fresnilloindicate a magmatic contribution is based ongenetic similarities found in a number of othernearby Ag- base-metal deposits which form ametallogenic belt that transects northern Mexico(Clark et al. 1982). Other than their containedmetals, these deposits share similar ages ofmineralization, close association with magmatism,and occurrences of saline fluid-inclusions. someof which, based on their enriched stable isotopes,very high salinities exceeding 35 eq. wt. o/o NaCl,and proximity to intrusions, are clearly magmaticin origin (e.9., Rye 1966: Sawkins 1964).Therefore, chloride in the Fresnillo brines seemsto be not only of magmatic origin, but wasnecessary to the introduction of metal-transportingf luids.

Orygen and Hydrogen IsotopesStable-isotope studies of epithermal deposits

are largely restricted to quartz, calcite, and

20,000

I contribution ofaq.

t carbon dioxidetoTm

Santo Nifro Vein425 Level West-Central

Stage l l Stage l l l Stage lV

* * s E ; * g ; * * so

llo

0

250"

O- 6= zE s' i ;

F }r

(J

I

EooE@

F

M E -

Figure 9. Sharp variations in fluid salinit ies over time during formation of the Santo Nino vein, Fresnil lo District;data compiled from observations of the west-central sector on the 425 level. Quartz and calcite volumetricallydominate the vein-fi l l ings of stages l, II, and III, for which silver sulfosalts and base-metal sulfides are the ore-bearing minerals; stage IV consists only of calcite and is barren. No data are available for sulfides in stage III. Brinepulses are associated with deposition of sulfides and sulfosalts in Stages I and II, whereas dilute fluids areassociated with deposition of quartz and calcite gangue. See Simmons et al. (1988) and Simmons (1991) for furtherdescr ipt ion and d iscussion.

466

"adularia" gangue, as these are the most commonoxygen-bearing minerals. The 6l8O.in.,ul compo-sitions are converted to 6lsowater compositionsfrom mineral-water equilibrium fractionationequat ions (e .9 . , F r iedman & O 'Ne i l 1977;Matsuhisa et al. 1979) and estimated equilibriumtemperatures, the latter typically determined fromfluid-inclusion homogenization data. The mini-mum range of homogenization temperatures forany one generation of primary fluid-inclusionsfrom these deposits is about 20 "9, correspondingto an uncertainty of about * 0.5

"/no in calculating

equilibrium water compositions. Note that oxygenisotopes can also be measured directly oninclusion fluids (e.g., Yigk et al. 1994), thoughsuch results are not available for the depositsdiscussed here.

Hydrogen-isotope compositions are measureddirectly on inclusion fluids released under vacuumby crushing; thus, ore-related sulfide minerals,along with quartz and calcite, can be analyzed.The uncertainties in bulk analysis are unknownbecause several fluid-inclusion generations maybe present, but average results suggest a range ofup 20

"/on are possible. The 6D data are difficult to

interpret without knowledge of the meteoric water

Low-s ulfi dat ion Ep i t her ma I Dep os i ts

Figure 10. The shaded regions representthe range of quartz oxygen-isotope compo-s i t ions versus equi l ibr ium temperalure.estimated from fluid-inclusion data, forepithermal deposits (Tables 3, 4). Thecurves represent calculated compositions oflocal meteoric (-6 to -18 "/*)

and magmaticwater (8

o/oo; in equilibrium with quartz as a

function of temperature, based onfractionation factors determined bvMatsuhisa et al. (1919\.

6t8o %n

composition at the time of mineralization. Thecomposition of local meteoric water may beinferred from trends in combined 6r80- 6D data;in very young deposits, modern meteoric watercompositions can be used.

Oxygen-isotope compositions of quartz fromthe six epithermal deposits are plotted versustemperature in Figure 10. Curves representing theisotopic compositions of magmatic (6t80 : 8

o/oo)

and local meteoric water in equilibrium withquartz are shown for comparison. The 6180values of fluids from these deposits either plotbetween meteoric and magmatic compositions(Antamok-Acupan, Comstock, Hishikari,Tayoltita) or overlap with the magmaticcomposition (Emperor, Fresnillo), with maximumpositive enrichments relative to local meteoricwater that range from 7

o/oo 6180 at Hishikari to

l8-20o/no 5l80 at Comstock, Emperor, andFresnillo. Accepting a value of about 9oloo 6180 a,the maximum possible for enrichment to meteoricwater due to water-rock interaction (as previouslydescribed), suggests that the waters at Comstock,Emperor, and Fresnillo had minimum magmaticwater inputs of 50o%, and the waters at Hishikari,Tayoltita, and Antamok-Acupan had no magmatic

Comstock

Tayoltita

Hishikari htrdn

Fresnillo,r{

Emperor

,y{

467

S.f-. Simmons

input but instead were derived from shifted

rneteoric water. Alternatively, assurning that there

were no positive enrichments due to meteoric

water-'rock interaction. the minimum magmatic

inputs into deposit waters ranged frorn about 100%

at I l ishikar i to 90yo at Emperor and Fresni l lo.-fhese two models bracket the range of possible

interpretations. 'Ihe results from Hishikari, which

indicate changes in the 6l80 water composit ion

over time. are discussed further below.

Matsuhisa & Aoki (1994) used the banded

quartz-adularia-sLrlfide vein fil l ing of the Ryosen

5 vein (average vein grade of 280 g Au/t)

Fl ishikar i to assess changes in 6'oo'u,, , . . compo-

si t ions with t ime. Highest ore grades within the

vein occur in vcry f ine-grained sulf ide-r ich

Ginguro bands that l ie near and paral lel to the

rvallrock contacts. Sulfide-poor bands are

conrposed of fine-grained quartz and ofthoclase

(adular ia) that are di f f icul t to separate physical ly.

By ingeniously comparing 6lsOqurnz-adurar ia versus

the quartz"-adularia ratio tlf the material analyzed

(determined frorn XRD and Al analyses),

Matsuhisa & Aoki (1994) def ined two dist inct

linear trends. which in turn were used to interpret

pure encl-member compositions ftrr 6'8Our"r,, and

E'tO"dul", iu associated with high-grade (>1000 g

Auit), and low-grade (<20 g Au/t) stages of vein

fill ing; late barren quartz was also analyzed'

Oxygen-isotope equi l ibr iurn temperatures were

calculated for quartz-adularia pairs, and

homogenization temperatures were measured for

barren quartz. The results indicate 6'tO*u,.,

composit ions of -0.1o/o., for f lu ids associated with

Au-Ag mineral izat ion, in contrast^to 618O*u,. ,

compositions of -3.6 and -4.3 to -5.7"/uo for fluids

associated with later low-grade and barren stages,

respectively. Thus a rninimum enrichment of

about 3.5"/oo 6'tO is indicated for "ore-stage"

waters relative to "gangue-stage" waters,

sLrggesting that mineralizing fluids contained

inputs of 30oh- or more. magmatic water.

Hydrogen- and oxygen-isotopic compositions

of waters associated with four of the epithermal

deposits (Table 4) are plotted in Figure 11, and

show broad patterns further suggesting that the

mineralizing fluids were mixtures of meteoric and

magrnat ic water (Ahmad el al . 1987; O'Nei l &

468

6"0 %.

Figure ll. fhe oxygen and hydrogen isotopiccompositions of epithermal fluids for epithermaldeposits (Tables 3, 4). The positions of local meteoricwater for Antamok-Acupan and Emperor are basedon present-day compositions, whereas fbr Comstockand Fresnillo thev are estimated fiom stable-isotopetrends.

Si lberman 1974; Sau'kins et al . 1979; Simmons e/

,z/ . 1988; Vikrc 1989). Boi l ing cannot account for

these trends as indicated by the calculated

enrichments due to adiabatic steam loss (Fig. 5c).

Here again the Fresnillo and Emperor data plot

close to the field of magmatic water, though their

deuterium values are difficult to distinguish from

the 6D compositions of local Ineteoric waters. A

much stronger indication of a magmatic compo-

nent in ore fluids is suggested by' the Comstock

data. based on two different lines of evidence: l)

the increase in the positive slope of trajectories

connecting meteoric and hydrothermal waters as a

function of increasing 6D, accepting that the

source of meteoric water remained constant

during the period of mineral izat ion (Vikre 1989),

and 2) the position of one sample associated with

bonanza silver ore from the 1200-foot level ofthe

Con Virginia mine, which has a water

composition of 1.9 u/,,,,

6'*o and -69 n/oo 6D

(O'Neil & Silberman 19'14). Accordingly, the

proportion of magmatic input into Comstock

fluids ranges from about 30 to 75 %o'

The wide range of 6D values indicated for

both Emperor and Cornstock data sets deserve

further comment in the light of their potential

relation to inputs of magmatic waters. Taylor

6 D % .

x Emperor

1 Frssni l lo

O Antamok-Acupan

. Comst@k

F

(gECE

$

III

i

(1986, 1992) has shown that the 6D composit ionof water vapor in equilibrium with a hydrousmagma decreases as the fraction of waterremaining in the melt decreases, and that the

resulting range of 5D water-vapor compositionsdepends on the mode by which a hydrous magma

degases. Thus, the 6D compositions of watervapor evolved through continous open-system andclosed-system degassing range from -25 to <-125n/uo

and -25 to -60 o/oo, respectively, as the fraction

of water remaining in the rnelt decreases from 1 to

0.1 (Taylor 1986, 1992). The 5D values ofEmperor fluids range from - l0 to -58o/oo,

overlapping with magmatic vapor composit ions

derived from either mode of degassing. The 5Dvalues for Comstock fluids^are much lighter andrange from -69 to -133"1oo, suggesting thatmagmatic vapors were mostly derived from open-system degassing. The relatively wide range of

6D compositions in these two deposits maytherefore relate to magma degassing processes.

Helium IsotopesHelium isotopes were measured for vein

materials obtained from subsurface mineworkings (>100 m depth) at Fresnillo andAntamok-Acupan (Table 4). Inclusion fluidshosted by quartz, calcite, and sulfides werereleased under vacuum by thermal decrepitationor crushing 10- to 20-g monomineral ic samples;their isotope compositions were measured on aNier-type double focussing mass spectrometer.The R/Ra values, between 6 to

'7 at Antamok-Acupan and between I and 2 at Fresnillo (Fig.12), indicate a component of mantle He,presumably transported to shallow crustal levelsby ascending magmas (Simmons 1986; Simmonset al. 1988).

The Fresnillo results seem low for arc-relatedfluids, and they may result from three differentprocesses: 1) preferential outward diffusion of lHe

over "He; 2) in-.situ radiogenic accumulation of*H"; or 3) radiogenic accumulation of -He

associated with long magma residence times andcrustal contamination. Of these, the last seemsmost likely, given the relatively thick continentalcrust through which Fresnillo magmas migrated(see Simmons et al . 1988). The hel ium-isotope

Low-su lfidat ion Ep it herma I De p o s i t s

O A Fresnil lo

a O Baguio

t A t Aatmo.ph"r ic h" l iur l

He concentration (x 10 6

) cc STP/g

Figure 12. Helium-isotope compositions (R/Ra)versus helium amount for inclusion fluids contained inquartz, calcite, and sulfides from Fresnillo, and quartzfrom Antamok-Acupan (Simmons 1986; Simmons eta / . 1 9 8 8 ) .

ratios of geothermal emanations associated witharc volcanism and similarly thick crust range from1.30 to 2.16 R/Ra in the Southern Volcanic Zone.andl.44 to 6.47 R/Ra in Central Volcanic Zone ofthe Andes, thus supporting this interpretation(Hilton et al. 1993). For Fresnillo, the resultsfurther indicate relatively uniform compositiousthrough time irrespective of mineral host or fluid-inclusion salinities, confirming that processesgoverning helium input are decoupled from thoseaffecting other fluid components (includingmetals), consistent with observations of TVZgeothermal fluids and data from otherhydrothermaldeposits (Simmons et al. 1987).

Note that investigators wishing to pursue He-isotope analyses of epithermal materials shouldensure that their samples are shielded from theeffects of cosmic radiation, especially at higherelevations of about 1500 m asl or more, ascosmogenic 'H" "un be produced by nuclearreactions involving spallation or neutronabsorpt ion by'Li (e.g ,Kurz 1986). Such effectsare interpreted for surface vein-quartz samples at2200 m asl from the Fresnillo district, which haveanomalous values exceeding 100 R/Ra (Simmons1986; Simmons et al . 1986). The observedexponential decrease in R/Ra with depth (from

r

S.F. Simmons

I l6 R/Ra at the surface to 65 R/Ra at l.l m depth)along with calculat ions of the cosmic-rayattenuation length in rocks, however, indicate that

these radiation effects are unlikely to penetrate

depths greater than about l0 m at Fresnillo(Simmons 1986; Simmons el al . 1986).

N2-Ar-He RatiosProblems associated with the analysis of gas

species in f' luid inclusions are related to loss of H2

and H2S (through diffusion and post-extraction

reaction; see Graney & Kesler this volume) and

these artefacts alter the redox state calculated for

publ ished data(e.g., Roedder 1984; Hedenquist er

al. 1992). However, such problems should not

affect measurement of N, Ar and He, though there

are few available measurements on epithermal

materials. Norman & Musgrave (1994) reported

data from three epithermal deposits, including the

Santo Nif lo vein, Fresni l lo (Fig. l3). Gases were

measured by a quadrupole mass spectrometer for

fluids released under vacuum by thermal decrepi-

tat ion or by crushing of 0.1 to 5 g of inclusion-

bearing material. The two smaller deposits from

New Mexico (not\shown) both have gas trends

indicating a possible "basaltic" signature. The

Fresnillo data (Benton 1991) form a broad linear

pattern that roughly overlaps with the mixing

envelope having andesitic and meteoric end-

members. These results are consistent with

salinity, and stable- and helium-isotope data,

which support the existence of magmatic

contributions in the Fresnillo fluids. Unfor-

tunately, data from coexisting fluid inclusions and

stable isotopes are unavailable to assess co-

variations or further constrain this interpretation'

DISCUSSION

In this paper I have attempted to document the

main geochemical evidence which indicates

magmatic contributions to low-sulfidation epi-

thermal environments. The strength of this

evidence is bolstered by studies of active systems'

where fluid compositions from geothermal

systems in different stages of evolution can be

compared on a regional and global scale. Besides

this temporal constraint, the capacity to analyze

all fluid components means that interpretations of

410

Fresni l lo

2000

N , / H e

1000

looo He 2oo Ar

Figure 13. Relative concentrations of nitrogen, argonand helium in gases from Fresnillo inclusion fluids(Benton 1991,Norman & Musgrave 1994).

one component's origin can be checked forinternal consistencies by comparing it to othercomponents that behave in a similar manner (e.g.,

oxygen and hydrogen isotopes can be compared tochloride, and helium isotopes can be compared toCOz, N2, and Ar). ln epithermal deposits,temporal constraints are relative as determinedfrom the mineralogic record, with millimeter tocentimeter-scale parageneses restricted todistances of a few hundred meters or less; hence,interpretation of spatial variations in thecompositions of paleo-fluids at a fixed point intime across a deposit is extremely difficult. Dataquality is also restricted by the errors inherent toanalyses and interpretations of minerals andinclusion fluids. Thus, active geothermal systemsprovide a scale of comparison for spatial and

temporal relations and a framework forinterpretation not available from study of low-sulfidation deposits.

With these caveats, helium isotopes and N2-Ar-He ratios can provide the most diagnosticevidence of magmatic contributions to epithermaldeposits; the few available data indicate that thesecomponents are promising tracers of fluid originsand deserve much further investigation. Incontrast, enrichment in both oxygen and hydrogenisotopes relative to meteoric water, and highconcentrations of chloride, can provide permissive

100Ai r

evidence of magmatic contributions (Table 1).Currently, Fresnillo is the only deposit for whichall of these techniques have been applied, thoughthere is a gap in the continuity of samplesinvestigated from this deposit, and the isotopestudies are reconnaissance in scale.

Although there is evidence of magmaticcontributions, it should be clear that the dominantsource of water entering most low-sulfidationepithermal environments is meteoric, but this isnot the issue here. Instead, a number ofresearchers have summarily discounted thepossibility of magmatic contributions in magma-related ore-forming hl,drothermal systems,arguing that water-rock interaction is sufficient toexplain enrichments both for oxygen and hydro-gen isotopes, and the origins of other components,including metals (e.g., Taylor 1973 Campbel l eral. 1984; Seal & Rye 1992), notwithstanding thefact that a few percent magmatic water could alsoaccount for the same isotopic enrichments (e.g.,O'Neil & Silberman 1974 Sawkins et al. 1979).The problem then relates to the framework ofinterpretation, and I believe this involvesappreciation: I ) of the nature of magmaticcomponents, 2) that magmatic componentspotentially contribute to ore formation, and 3) thatmagmatic contributions can reach shallow epi-thermal environments.

The nature of magmatic components is bestunderstood from examination of degassingvolcanoes and study of porphyry ore deposits(".9., Hedenquist & Lowenstern 1994). Thesecomponents are mostly volatile and include water,carbon dioxide, chlorine (as HCI), sulfur (as SO2and H2S), and base and precious metals, all ofwhich are observed in low-sulfidation epithermalenvironments, with chlorine and sulfur beingimportant for metal transport. The signatureswhich record the appearance of magmaticcomponents in the epithermal environment aremostly restricted to those in Table l. Otherpotential tracers, such as the isotopic compo-sitions of carbon, sulfur, and lead, are commonlyambiguous due to effects relating to redox state orcrustal contamination, and are difficult to interpret(see Hedenquist & Lowenstern 1994).

That magmatic fluids can reach and influencelow-sulfidation epithermal environments is

Low-suffidation Epithermal Depos its

probably better documented than mosr geo_scientists realize, with a much clearer magmaticconnection existing for high-sulfidation epi-thermal environments (Arribas this volume). Atone extreme are the eruptions of magma throughgeothermal systems, which in recent historyinclude the 1886 eruption of Mt. Tarawera in NewZealand (Simmons el a/. 1993), the 19?6-1977eruption at Krafla in Iceland (with the firstrecorded discharge of magma from a geothermalwell: Larsen et al. 1979), and the l99l eruption ofMt. Pinatubo, Philippines, formerly a geothermalprospect of the Phi l ip ine Nat ional Oi l Company(with two pre-eruption exploration wells: Delfin etal. 1992). At the other extreme is the evidence forsteady influx of magmatic components (helium,nitrogen, chlorine, and water) into geothermalsystems and epithermal deposits. These extremesalso represent end-members on a time-scale ofinfluence, one nearly instantaneous, from hours todays, and the other continuous, over hundreds totens of thousands of years. At time-scales inbetween are the pulses of fluid that reflect themagmatic inputs inferred for Fresnillo, Hishikari,and Comstock. These sharp changes in fluidcompositions have not been recognized in activesystems, though wells have only been monitoredfor a maximum of about 35 years (e.g., Wairakei).Fluid pulses are also known from the mineralogicrecord of some active geothermal systems; forexample, 618O.ut.it" values at Kawerau indicate theformer field-wide presence of a carbon dioxideand 6r8o-enriched thermal fluid (up to 5

"/un

compared to current values of -3.75 '' ln,,) of likely

magmatic or igin (Christenson 1989). Thus bothtransient and persistent influxes of maglnaticcontributions are possible.

Only at Fresni l lo, for which high-sal ini tybrines of rnagmatic origin are interpreted, can acause and effect relationship between magmaticinputs and mineralization be considered. For otherdeposits, such as Hishikari and Comstock. theavailability of metal-transporting ligands inmineralizing fluids cannot be assessed by currentanalytical techniques and, therefore, any geneticlink between magmatic inputs and mineralizationis inferred only by spatial association betweenprecious-metal occurrence and isotopicallv (6180

S.F. Simmons

and 6D) enr iched gangue minerals . Even in the'lVZ.

where relatively high concentrations of

precious metals are being fluxed in "gassy"

geothermal l lu ids conta in ing magmat ic

contributions, the cause and effect relation is

ambiguous as the source of aqueous sulfur,

assuming it accounts for the aqueous Au and Ag,

cannot be traced. The ultimate source of metals in

low-sulfidation epithermal deposits thus remains

poorly understood.

ACKNOWLEDGMENTS

I thank J. W. Hedenquist , J . Margol is , R.

Sher lock, and J. F. H. Thompson fbr thei r per-

ceptive comments and crit icisms of an earlier

version of this paper, and thank Louise Cotterall,

who drafted some of the Figures.

REFERENCES

AHMAD, M., SOLOMON, M. & WALSHE, J. L.

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