Lithocap Hedenquist2013.Full

Bulletin of The Society of Economic Geologists

Vol. 108

November

No. 7

Modeling the Formation of Advanced Argillic Lithocaps: Volcanic Vapor Condensation Above Porphyry Intrusions Jeffrey W. Hedenquist1,† and Yuri A. Taran2 1 Department 2 Institute

of Earth Sciences, University of Ottawa, Ottawa, Ontario, Canada K1N 6N5

of Geophysics, Universidad Nacional Autónoma de México, Coyoacán, México D.F. 04510, Mexico

Abstract Hypogene advanced argillic alteration, typically quartz-alunite with halos of kaolinite ± dickite and roots of pyrophyllite ± diaspore, forms in the epithermal environment from condensates of magmatic vapor that contain SO2 and HCl, which exsolved from an underlying intrusive source. The most aggressive, nearly isochemical leaching of the host rock by the most acidic condensate, commonly pH ~1, leaves residual silica that recrystallizes to quartz within the flow channel at a high condensate/rock ratio, forming the core of alteration. The alteration characteristically flares upward along feeder structures, and if a permeable lithologic unit is intersected, the alteration zones mushroom to form a subhorizontal blanket due to lateral flow. Where subsequently mineralized, the residual quartz, commonly with a vuggy texture that reflects the texture of the original lithology, has higher Au (and Cu) grades than the quartz-alunite halo. Tonnage in these high sulfidation systems may develop within the subhorizontal lithocap, although the highest grades are typically confined to the structurally controlled feeders. We modeled a typical volcanic vapor condensate, starting with the measured composition of ≤877°C fumaroles from Satsuma Iwojima rhyolite dome, Japan, as it cooled and reacted with a host rhyolite; the predicted hydrothermal mineralogy reproduces the alteration pattern observed in lithocaps that host high sulfidation deposits. The modeling confirms that aluminum-rich minerals (pyrophyllite, diaspore, locally andalusite) are stable at higher temperature at depth, whereas at lower temperature and shallower depth, Na and K alunite become stable. At the lowest temperature (10:1), and where SO2 >>H2S in the original volcanic vapor; this SO2-dominant composition is typical of andesitic to rhyolitic volcanoes. The reason for this mineral transition, and the upward flare (widening) of the alteration zone along structures, is related to the dissociation and increased reactivity of H2SO4 and HCl as the temperature decreases. Below ~200°C, only quartz, pyrite, native S, and anhydrite are stable, hence the formation of the dominant quartz from the silica residue. A further check on our modeling is the observation that the calculated composition of the condensate after reaction with fresh Satsuma Iwojima rhyolite and alteration minerals, and cooling to 100°C, is similar (within a factor of two) to that of acidic springs, with pH ~1, that discharge around Satsuma Iwojima and other active volcanoes. The most extensive lithocap alteration, residual quartz and/or quartz-alunite, is commonly offset from the surface projection of the causative intrusion. This observation can be explained by a combination of two factors—hydrology and temperature. Due to hydraulic gradients at shallow depths in a volcanic edifice, the acidic condensate tends to flow along permeable lithologic units away from the locus of the high-temperature vapor plume, which rises directly over the intrusion. Where lateral flow occurs, the most intense leaching and widespread advanced argillic alteration, which develops largely at temperatures of 20 km2; Sillitoe, 1995) and, since they are resistant to erosion, are typically prominent at surface, which can make them easy to recognize. Despite the relative ease of identifying lithocaps, their large size means that it is commonly difficult to identify the position of the underlying causative intrusion(s). Here we review hypogene advanced argillic alteration (Meyer and Hemley, 1967), as well as the magmatic source of acidic fluids that is responsible for this alteration type. An early understanding of the conditions of formation of advanced argillic alteration came from experimental studies, for example, the high-temperature stability of pyrophyllite, diaspore, and andalusite (Hemley et al., 1980; Sverjensky et al., 1991), and the stability of alunite (Hemley et al., 1969; Stoffregen et al., 2000) as well as that of other minerals in this assemblage.

NW

We then use the measured composition of magmatic vapor discharged from a rhyolite dome and its acidic condensate, after reaction with rock, to check the results of reaction modeling of the fluid composition as well as mineralogy and temperature stability (i.e., zonation) of the alteration products that occur. We find that the results are sensitive to the initial redox state of the magmatic fluid, the condensate/rock ratio, and the temperature of condensate-rock reaction. Our model results provide insight into the nature and patterns of advanced argillic halos to residual quartz alteration, for example, why residual quartz (recrystallized from the strongly leached siliceous residue of rocks) and quartz-alunite flare upward along structures (thus apparently having no root zones). The results also reveal why the main body of residual quartz (locally vuggy in texture, reflecting the original rock texture) may be offset from the underlying intrusive source, locally for distances of up to a few kilometers, rather than forming directly above it (Fig. 1). This potential for offset of the largest portion and most strongly altered cores of lithocaps from the causative intrusions is essential to appreciate during exploration for and assessment of both epithermal and porphyry deposits. Advanced Argillic Lithocaps and Porphyry Deposits Shallow intrusions are the source of magmatic volatiles that ascend and discharge as volcanic fumaroles (e.g., Giggenbach

SE

Lepanto highsulfidation deposit

elev. m 1400

Imbanguila dacite

Bato dacite residual quartz

kaolinite

1000

enargite-Au

900 mL qtz-alunite

2.5 % Cu eq.

qtz-alunite

1.4 Ma

500 m 600

Far Southeast porphyry

.

breccia pipe

1w

t%

Quartz-alunite

Basement

eq

Residual quartz (silicic)

chlorite ±illite 1.35 Ma

Cu

Enargite-Au ore

pyr op

hyl

lite

dickite ± diaspore

Qtz diorite potassic biotite 1.4 Ma

Fig. 1. Schematic NW-SE long section through the Lepanto lithocap and its enargite Au orebody, offset to the NW from the underlying Far Southeast porphyry deposit (Garcia, 1991). Long dashes show the unconformity along which the lithocap is focused, where it is intersected by the Lepanto fault (in the plane of the long section). The enargite Au ore is largely hosted by residual quartz, restricted to within ~2 km of the intrusive source of volatiles; the residual, vuggy quartz has a halo of quartz-alunite and kaolinite ± dickite alteration, which continues past the residual quartz along the unconformity for a additional 2 km to the NW (Chang et al., 2011). The paleosurface over the porphyry was ~1,500- to 2,000-m elevation at the time of intrusion at ~1.45 Ma (Hedenquist et al., 1998). Biotite associated with potassic alteration in the porphyry deposit and alunite in the lithocap are contemporaneous at ~1.4 Ma (Arribas et al., 1995); these zones were later overprinted by white mica alteration (1.3−1.35 Ma), which is transitional upward to pyrophyllite (Hedenquist et al., 1998). If the lithocap were exposed by erosion to ~900-m elevation (900-mL drift shown), drilling near the vertical shaft, i.e., where the residual quartz is thickest, would not intersect the causative intrusion, its porphyry deposit, or even proximal porphyry-related alteration.

200



MODELING OF HYPOGENE ADVANCED ARGILLIC ALTERATION BY VOLCANIC VAPOR CONDENSATES

et al., 1987); a portion of the vapor condenses (Hedenquist et al., 1994) to form hypogene advanced argillic alteration (Ransome, 1907; Hemley et al., 1969, 1980; White, 1991; Hedenquist et al., 1998, and references therein). This advanced argillic alteration typically overlies porphyry deposits, where not eroded, and can also host epithermal mineralization (Sillitoe, 1999, 2010). Since both porphyry and epithermal deposits tend to form proximal to their causative intrusions (Sillitoe, 1999, 2010), they reflect the intrusive source of fluids that form characteristic and zoned alteration and mineralization (Einaudi et al., 2003; Sillitoe and Hedenquist, 2003). It is common for the altered rocks that host ore within the epithermal environment to be offset from, and thus asymmetric relative to the surface projection of the parent intrusion (Sillitoe, 1999). A good illustration of this is the Lepanto lithocap and its contained high sulfidation Au-Cu deposit, Mankayan district, Philippines, which are offset up to 4 and 2 km, respectively, from the quartz diorite dikes that acted as the fluid source (Hedenquist et al., 1998) and formed the Far Southeast porphyry Cu-Au deposit at depth (Fig. 1; Garcia, 1991; Chang et al., 2011). The reasons for the potential offset of lithocaps from their underlying intrusions are many and include geologic heterogeneity that in turn may influence the hydraulic gradient in the near-surface, epithermal environment. Intrusive centers associated with volcanic activity had volcanic edifices over their tops, prior to erosion, which would have resulted in hydraulic gradients down the slopes of the volcanic edifices (Sillitoe, 1994, 1999, and references therein). There are two characteristic features of the geometry of residual quartz and its advanced argillic halo, the latter dominated by quartz-alunite that grades outward to kaolinite ± dickite, which may be transitional at greater depth to pyrophyllite ± diaspore. First, Steven and Ratté (1960) noted that these characteristically zoned alteration zones at the Summitville deposit, Colorado, flare upward along structural and hydrothermal breccia conduits, with the residual quartz cores being the principal hosts to subsequent enargite and Au mineralization. Secondly, Sillitoe (1995, 1999) observed that where such structural conduits intersect permeable lithologic horizons or unconformities, such as those at Yanacocha, Peru, or Lepanto, respectively, the structurally controlled, ascendant fluids will flow laterally along these lithologic horizons. Source of Hypogene Acidity Volcanic vapor and fumarolic discharges On exsolution from magma, an aqueous fluid commonly consists of a single phase of intermediate salinity and density (Shinohara, 1994, 2009; Audétat et al., 2008, and references therein). If the fluid intersects its solvus on ascent and depressurization, it will separate into low-salinity vapor and dense, hypersaline liquid (Sourirajan and Kennedy, 1962; Henley and McNabb, 1978; Shinohara, 1994; Heinrich, 2007). Where the salinity of the exsolved single-phase intermediate density fluid is low, typically 900°C) fumaroles from a variety of quietly erupting to passively degassing volcanoes is well known, with useful compilations including those by Giggenbach (1987, 1996), Symonds et al. (1990, 1994), Taran et al. (1995, 2001), Hedenquist (1995), Einaudi et al. (2003), and Fischer (2008). The vapors from felsic to mafic volcanoes have a wide range in composition; the major component is H2O, typically from ~87 to 97 mol %, followed, in decreasing amounts, by CO2 (1−6 mol %, although lower and higher values also occur), SO2 (0.5−3, locally to 7 mol %), HCl (0.1−1+ mol %), and H2S (0.1−1 mole %), with traces of N2 and other gases; HF is commonly an order of magnitude less than HCl, mainly 600°C, vapors discharged at the surface from the condensed magmatic-hydrothermal systems may have temperatures up to 400°C (Einaudi et al., 2003, and references therein) and occur beneath summit craters, adjacent to high-temperature volcanic fumaroles, or on the flanks of the volcanoes. Reyes et al. (1993) referred to the transition between the hightemperature vapor zone and the magmatic-hydrothermal system at Alto Peak, Philippines, as a vapor-cored “chimney” enveloped by the adjacent condensed system. This vapor chimney acts like a “bellows” (Giggenbach, 1975, 1982, 1987), inflating (on heating) or deflating in size, or even closing,

depending on the temperature and intensity (heat flux) of the magmatic discharge. In addition to chemical indications (Giggenbach, 1987), isotopic evidence confirms that these reactive fluids originate by condensation of magmatic vapors into meteoric water (Giggenbach et al., 1990; Hedenquist et al., 1994; Taran et al., 1995, 1996). The magmatic-hydrothermal liquids may discharge at the surface as hot springs with a pH as low as ~1.0. For example, at Satsuma Iwojima, Japan, and Ebeko and Baransky, Kuril Islands, the measured pH of springs ranges from 1.1 to 1.7, with the most acidic spring having the largest proportion of magmatic vapor condensed into the local meteoric water (Hedenquist et al., 1994; Taran et al.,



MODELING OF HYPOGENE ADVANCED ARGILLIC ALTERATION BY VOLCANIC VAPOR CONDENSATES

1996). Volcanic crater lakes, such as Kawah Ijen in Indonesia, Kusatsu Shirane in Japan, Poás in Costa Rica, and Ruapehu in New Zealand, or boiling pools in volcanic craters such as Black Pool at Mutnovsky volcano, Kamchatka, can have even lower pH values, some negative, due to the effects of evaporative concentration on H2SO4 and HCl in the hot lakes, after the fumarolic vapors are condensed into the crater lake waters (Rowe et al., 1992; Taran et al., 1992; Delmelle and Bernard, 1994; Christenson, 2000). The redox state of magmatic-hydrothermal vapors is controlled by the sulfur-gas buffer at temperatures >350°C (Giggenbach, 1987; Einaudi et al., 2003; Zelensky and Taran, 2011), but compositions are more reduced and H2S dominant at lower temperatures (Fig. 2). Condensation and formation of the aqueous liquid phase leads to the redox control being provided by the dissolved sulfur species (HS–/HSO4– or H2S(aq)/ HSO4–). Both gas and aqueous S buffers control redox state at high fluid/rock ratio, whereas at lower temperature and lower fluid/rock ratio, subsequent reaction with Fe-bearing minerals (the rock buffer) causes reduction of the fluid, as reflected in the vapor discharged from magmatic-hydrothermal systems (Fig. 2; Giggenbach, 1987; Einaudi et al., 2003). The dominant chemical process in condensed magmatic vapors that are cooling is aqueous redox-disproportionation of the absorbed SO2 according to the reactions (Iwasaki and Ozawa, 1960; Sakai and Matsubaya, 1977; Kusakabe and Komoda, 1992):    4SO2(g) + 4H2O ⇔ 4H2SO3(aq) ⇔ H2S(aq) + 3H2SO4(aq) (1a) 3SO2(g) + 2H2O ⇔ S° + 2H2SO4(aq), (1b) followed by dissociation of H2SO4, which yields H+ ion and ionized aqueous-sulfate species: H2SO4(aq) ⇔ HSO4– + H+ (2a) HSO4– ⇔ SO42– + H+. (2b) H2SO4(aq) becomes a strong acid as temperature decreases, with the dissociation constant K2a at 100°C close to 102. However, at higher temperature, H2SO4(aq) becomes the predominant sulfate-bearing form (K2a ~0.01 at 350°C; Oscarson et al., 1988; Hnedkovsky et al., 2005). The second dissociation constant, K2b is >H2S), the latter close to the actual composition of the Satsuma Iwojima vapor (Table 1). Variations of this ratio in high-temperature magmatic fluids will result in different evolution pathways for the postmagmatic condensate-rock interaction. Thermodynamic data Thermodynamic data for gas and aqueous species plus minerals are compiled in the thermodynamic database of the Table 1. Chemical Compositions of the Satsuma Iwojima Rhyolite and High-Temperature Volcanic Vapor Used in the Model

(3)

At high temperature, equilibrium (3) is shifted to the left, leading to an increase in pH, as is also the case for H2SO4(aq) (e.g., Simonson and Palmer, 1993; Tagirov et al., 1997). Calculation Procedure Method Reaction modeling was performed with the HSC-7 computer code (A. Roine, 2007, HSC Chemistry, www.outotec. com), using the programs GIBBS and/or SOLVGASMIX for

1527

Rock composition Oxide

Wt %

Vapor composition Component

Mol %

SiO2 75.96 H2O 94.436 Al2O3 12.54 H2  0.678 FeO  2.14 CO2  2.544 Fe2O3   0.96 H2S 0.17 CaO  3.10 SO2  2.544 Na2O  4.13 HCl  0.594 K2O  1.18 N2  0.034 Notes: Data from Shinohara et al. (1993) and Saito et al. (2002)

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HEDENQUIST AND TARAN

commercially available HSC package, which was last updated for the HCS-7 version in 2009. Thermodynamic data for most aqueous species and minerals in HSC are taken mainly from SUPCRT92 (Johnson et al., 1992), updated to SUPCRT95 (Sverjensky et al., 1997); data for some aqueous species and minerals are compiled from other sources, up to 2005 (Table 2). Test calculations for different compositions of aqueous solutions using HSC-7 and SOLVEQ code (Reed and Spycher, 1984) with the updated SOLTHERM thermodynamic data set (mainly from the same SUPCRT95 base) provided essentially identical equilibrium compositions. Some of the data for SO42–-bearing species in the thermodynamic base were replaced with data obtained from more recent experiments, using the HSC option of “own data base” and a procedure providing internal consistency. We included data for H2SO4(aq), absent in the SUPCRT95 and SOLTHERM bases, using data of Hnedkovsky et al. (2005); data for the AlSO4+ complex were taken from Tagirov and Schott (2001). We recalculated erroneous thermodynamic parameters for SO2(aq) in the HSC base using data from Schulte et al. (2001), and data for NaSO4– were recalculated using experimental results from Pokrovski et al. (1995); those for NaHSO4(aq) and CaHSO4+ were included from the compilation by Kraynov et al. (2004). Data for HCl(aq) were fitted to the dissociation constant from Tagirov et al. (1997). Table 2 summarizes the

original data sources for all gaseous and aqueous species plus minerals used. Assumptions and limitations Chemical equilibrium between fluids and alteration minerals is a common assumption in geochemical studies of hydrothermal systems (Giggenbach, 1980, 1981, 1988; Reed, 1982, 1997; Reed and Spycher, 1984). In spite of the significantly different kinetics of dissolution and precipitation among various minerals, computed alteration assemblages from other studies commonly show a close similarity to natural alteration assemblages, even at temperatures as low as 25°C (e.g., Palandri and Reed, 2004; Ramirez et al., 2004; Heinrich, 2005). The choice of fluid-rock interaction model (closed or open system) is based on the principal objective of the modeling. The main question to be answered in this study is how and why the advanced argillic alteration mineralogy and its zoning, associated with a degassing magmatic body is influenced by the geochemical evolution of the magmatic vapor condensate as a function of the condensate/rock ratio and temperature. The acidity of the reactive liquid will vary with temperature because of the temperature dependence of the dissociation constants of the main H+-bearing components, H2SO4(aq) and HCl(aq); it is also related to the amount of rock components that are dissolved.

Table 2. Aqueous Species, Minerals, and Gases of the System Used in This Study, Plus Sources of Thermodynamic Data Aqueous species Aqueous species Al3+ Tagirov and Schott (2001) SO32– Al(OH)3 Tagirov and Schott (2001) SO42– AlOH2+ Tagirov and Schott (2001) Ca2+ Al(OH)2+ Tagirov and Schott (2001) CaCl2 Al(OH)4– Tagirov and Schott (2001) CaCl+ + AlSO4 Tagirov and Schott (2001) CaSO4 Cl– Sverjensky et al. (1997) CaHSO4+ Fe3+ Shock et al. (1997) Gases Fe2+ Shock and Helgeson (1988) FeCl2+ Sverjensky et al. (1997) H2O FeCl+ Sverjensky et al. (1997) H2 FeCl2 Sverjensky et al. (1997) SO2 + FeOH Shock et al. (1997) H2S + FeSO4 Bailey et al. (1982) S2 H2 Shock et al. (1997) HCl H+ Sverjensky et al. (1997) Minerals HCl Tagirov et al. (1997) H2S Plyasunov and Shock (2001) Al2SiO5, andalusite (and) HS– Shock and Helgeson (1988) Al2Si4O10(OH)2, pyrophyllite (pyr) HSO3– Shock and Helgeson (1988) FeO (FeO mineral proxy) H2SO4 Hnedkovsky et al. (2005) Fe2O3, hematite (hem) – HSO4 Shock and Helgeson (1988) Al2Si2O5(OH)2, kaolinite (kao) SiO2 Sverjensky et al. (1997) KAl3 (SO4)2(OH)6, alunite (alu) K+ Sverjensky et al. (1997) KAlSi3O8, microcline (mic) KHSO4 Sverjensky et al. (1997) NaAl3(SO4)2(OH)6, Na-alunite – KSO4 Sverjensky et al. (1997) NaAlSi3O8, albite (alb) KCl Sverjensky et al. (1997) SiO2, quartz (qtz) Na+ Sverjensky et al. (1997) FeS2, pyrite (py) NaCl Sverjensky et al. (1997) FeS, pyrrhotite (po) NaOH Sverjensky et al. (1997) Fe3O4, magnetite (mt) NaSO4– Pokrovski et al.,1995) S, native sulfur (S) NaHSO4 Kraynov et al. (2004) CaSO4, anhydrite (anh) – OH Sverjensky et al. (1977) KAl3Si3O10(OH)2, muscovite (mus) S2− Bailey et al. (1982) Ca2Al2Si3O10(OH)2, prehnite (pre) SO2 Schulte et al. (2001)

Shock et al. (1997) Shock and Helgeson (1988 Sverjensky et al. (1997) Sverjensky et al. (1997) Sverjensky et al. (1997) Sverjensky et al. (1997) Kraynov et al. (2004 Thermodata (1999 Thermodata (1999 Thermodata 1999 Thermodata (1999 Thermodata (1999 Thermodata (1999 JANAF (1998) Robie et al. (1979) Thermodata (1999) Thermodata (1999) Johnson et al. (1992) Stoffregen et al. (2000) Johnson et al. (1992) Stoffregen et al. (2000) Johnson et al. (1992) Johnson et al. (1992) Johnson et al. (1992) Johnson et al. (1992) Johnson et al. (1992) Thermodata (1999) Thermodata (1999) Barin et al. (1989) Shock et al. (1997)

Notes: Alunite and natroalunite are considered as individual minerals and not a solid solution; sources as cited in the HSC data base or added to the data base (latter in bold font)



MODELING OF HYPOGENE ADVANCED ARGILLIC ALTERATION BY VOLCANIC VAPOR CONDENSATES

To recognize these variations, it is sufficient to model a closed system that assumes fluid-rock equilibrium at different condensate/rock ratios and temperatures. In our calculations, we suppress boiling and formation of a coexisting vapor phase with the condensate by using a relatively high pressure of 200 bars, thus avoiding shallow boiling and vapor-rock interaction. Note that for a closed-system approach, all elements in the fluid-plus rock system are conserved. Therefore, the obtained equilibrium compositions of the aqueous phase may be compared to natural solutions only qualitatively (see discussion below). Another assumption is that the set of hydrothermal minerals chosen encompasses all those stable at low H+ activities (high pH), i.e., at low condensate/rock ratios. We use albite, microcline and muscovite, and prehnite as mineral proxies for the sinks of Na, K, and Ca, respectively, at low condensate/rock ratios. Another critical limitation is the valid range of thermodynamic data for some minerals and species. For most of the minerals considered here, their thermodynamic data are extrapolated to higher temperatures. However, data for key species that provide the acid-base balance and the pH of the aqueous liquid (SO2(aq), HCl(aq), H2SO4(aq)) are well known up to 350°C (Tagirov et al.,1997; Schulte et al., 2001; Hnedkovsky et al., 2005). The specific Gibbs free energies and equilibrium constants for the aqueous species correspond to the boiling-point curve (liquidvapor saturation). At high liquid densities (at 0.59 g*cm−3, the density of pure liquid H2O; thus, various thermodynamic issues associated with low densities (say, >H2S, the sulfate species are higher; at w/r >2, the associated sulfate species (H2SO4°, NaHSO4°) are more abundant as temperature increases. The compositions at 100°C (Fig. 3b) may model the character of thermal waters associated with either neutral pH, “mature” (Giggenbach, 1988) hydrothermal systems (where w/r = 2 and SO2 = H2S), or acidic Cl-SO4 springs and crater lakes (with w/r >2 and at higher SO2/H2S ratios). For a more accurate comparison, the modeling of acidic springs should include an increasingly complicated procedure involving a quasiopen system, since sulfur from the condensed volcanic vapor is subsequently lost from solution as it combines with Fe (pyrite), Al (alunites), and Ca (anhydrite) in higher temperature zones. Nevertheless, in general our modeled solutions show a common pattern of increasing concentrations of Al and Fe with decreasing pH; the Al concentration (Al3+ and other Al species) also shows a clear increase with decreasing temperature. The model behavior of Fe and Ca species is more complicated and reflects the effect of the closed-system approach. As a result, the calculated concentrations of both Ca and Fe are somewhat lower than usually observed in acid springs and crater lakes (e.g., Taran et al., 1992; Delmelle and Bernard, 1994), although the difference between observed and model values, starting with SO2 >>H2S, is minimal (Table 3).

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HEDENQUIST AND TARAN

w/r=2; SO2=H2S

1

Na

-

Cl

0.1

+

0.01

o

0.01 1E-3

NaHSO4

0

HSO4

250

0.1

o

NaCl

K

+ o

KCl

1E-3

o

H2S

o

SO2 NaHSO4

HSO4

3+

AlSO +

+

H2SO4

200

250

1E-3

-

NaSO4

o

+

1E-3 1E-4 100

Ca

2+

H2S

o

KCl

150

200

250

+

200

250

+

FeCl

300

350

w/r=10; SO2>>H2S -

HSO4

o

H2SO 4

-

Na

o

HCl

+

o

NaCl

0.01

o

SO2

+

FeCl

K

1E-3

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o

0

150

HSO4

AlSO4

CaCl2

o

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o

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-

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CaCl

0.1

H2S

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KCl

NaHSO4

o NaCl H SOo NaHSO4 4 2

0.01

Ca

AlSO4

2+

o

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Cl

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3+

Fe

FeCl2

o

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+

o

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1E-4 350 100

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150

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w/r=5; SO2>>H2S

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o +

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1E-4 100 350

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Cl Na 2+ Ca

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w/r=5; SO2=H2S

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KHSO4 KClo

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HSO4

+

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1E-3

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Aqueous species (mol/kg)

Cl

0.1 HS

2-

1E-4 100

Na

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o

NaCl o H2S

w/r=2; SO2>>H2S

1

Ca

3+

Al

2+

Fe

2+

SO4 +

1E-4 100 150 300 350 o Temperature ( C)

o

H2S

FeCl

200

+

2+

KCl

o

+

AlSO4 +

CaCl 0 KHSO4

250

300

350

Fig. 3. (a). Concentrations of all dissolved species (mole/kg, or molality) as functions of different condensate/rock (i.e., water/rock [w/r] of 2:1, 5:1, and 10:1) and SO2/H2S ratios in the condensed volcanic vapor, as the condensate cools and reacts with the rock. Note that concentrations of neutral nondissociated species (H2SO4°, HCl°) decrease sharply as temperature decreases, as these species increasingly dissociate. (b). Concentrations of the main solutes (mg/kg) and pH as functions of temperature for different condensate/rock (w/r) and SO2/H2S ratios in the condensed volcanic vapor. Note that for SO2 >>H2S, the resulting aqueous solution is sulfate dominant, whereas at SO2 = H2S, the Cl (mainly NaCl) component is dominant; the pH is always higher for the SO2 = H2S case. As the w/r ratio increases, the pH decreases from >3 at 350°C to >H2S, except for quartz, native sulfur, pyrite, and anhydrite, which are stable over the full temperature range of calculation. In the case of the H2S-rich

condensate at temperatures >H2S

pH 100000

K

100

Fe

Na

4

Fe 200

4

10000 1000

3

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SO4

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o

Temperature ( C) Fig. 3. (Cont.)

and lacks the K metasomatic minerals. At higher w/r ratios, the equilibrium mineral assemblage is similar for both H2Sand SO2-rich fluids but with lesser alumino-silicate minerals and more abundant alunites under more oxidized conditions. Pyrophyllite is increasingly unstable at w/r = 5, and kaolinite at w/r = 10, reflected by an increase of Al species (Fig. 3a) and total Al (Fig. 3b) in solution; diaspore and andalusite have a strong temperature dependence and become unstable at lower temperature, particularly andalusite, and diaspore stability is favored under reduced conditions. However, both Naand K-rich alunite are stable and dominate in abundance over the alumino-silicate minerals under more oxidized conditions. This reflects the increased abundance of available sulfate in solution and the fixation of both Na+ and K+ in this temperature range (Fig. 3b) at a w/r ratio >2:1. As temperature decreases below ~200°C, particularly in oxidized conditions and where condensate/rock ratios are high (>5:1), both Na and K alunite begin to dissolve (Fig.

4) due to the increased reactivity (decrease in pH caused by acid dissociation), accompanied by an increase in the Na and K species in solution (Fig. 3a). In the H2S-rich case and at w/r = 2, Na+ and K+ concentrations in solution are controlled by ion exchange between albite and microcline. Continued cooling at high condensate/rock ratios leads to the situation where only quartz (or its amorphous equivalent) and pyrite are stable, along with native S and anhydrite (Fig. 4); dissolution of all other minerals is reflected by the high Al3+, Na+, and K+ plus other species in solution (Fig. 3a, b). The dissolution of alunite—typically the major mineral component of advanced argillic alteration, after quartz—is primarily due to the increased stability of Al3+ and aqueous Al complexes at pH >H2S

w/r=2; SO2=H2S qtz

100

Mineral (g)

dsp

anh

S

10

alb

150

200

S

and

1

pyr 300 350

250

he

kao

100

Mineral (g)

qtz

150

200

250

pyr 300

350

w/r=5; SO2>>H2S qtz

S

dsp

anh

10

py

alu

alu

anh

Na-alu

S

py

Na-alu 1

1

and

kao

and

dsp

kao

0.1 100

150

200

250

0.1 350 100

300

w/r=10; SO2=H2S

100 S

Mineral (g)

and

100

10

anh

dsp

w/r=5; SO2=H2S 100

Na-alu

alu

10

kao

100

py

py

mus

mic

1

qtz

100

100

anh

py

Na-alu

1

and 200

250

qtz dsp

150

200

250

300

10

350

S

anh

1

alu Na-alu

py dsp and

alu 350

300

w/r=10; SO2>>H2S

qtz

10

0.1 100

150

0.1 100

150

200

250

300

350

o

Temperature ( C) Fig. 4. Amount of minerals (in grams) precipitated from the cooling of 1 kg of aqueous solution (condensate) as a function of temperature at different condensate/rock (w/r) and SO2/H2S ratios (the latter taken from the composition of the volcanic vapor that condenses). As w/r and SO2/H2S ratios increase, diaspore, andalusite, and pyrophyllite begin to form over primary minerals (muscovite and feldspars) at high temperature; at increasing w/r and SO2/H2S ratios and decreasing temperature, quartz, alunite, and other advanced argillic minerals plus pyrite are stable. Dissolution of Al from the rock (Fig. 3a) at lower temperature starts to leave a siliceous residue plus alunite and natroalunite, native S, pyrite, and a sulfate mineral (anhydrite; see text); at >H2S, leaving mainly residual quartz. Abbreviations listed in Table 2.

concentration in solution is controlled by small variations in the amount of pyrite, with its minimum corresponding to the maximum dissolved Fe (Figs. 3b, 4). Concentrations of Ca2+ and other Ca species are controlled by the presence of anhydrite and the availability of SO42– ion that, in turn, depends on the stabilities of the associated SO4 species (H2SO4°, HSO4–, NaSO4–, CaSO4°, etc.). The solubility of anhydrite increases with decreasing temperature (e.g., Corti and FernandezPrini, 1984), and the associated sulfate species become less

stable; this is why total Ca concentrations in solution generally increase on cooling (Fig. 3b). However, in the SO2-rich case, Ca concentrations are relatively low but roughly constant, reflecting the overall higher concentrations of SO4-bearing species. Effects of the condensate/rock ratio The titration of the acidic condensate by an increasing amount of fresh rock, i.e., starting from a condensate-dominated



MODELING OF HYPOGENE ADVANCED ARGILLIC ALTERATION BY VOLCANIC VAPOR CONDENSATES

SO2=H2S

SO2>> H2S o

o

100 C alb

100

hem 1

kao

pyr

0.1

MIneral (gram)

anh dsp alu py

pre

mt

mus

mic

qtz

mus

100 C

alb 100

S

mic

10

10

alu qtz

S

hem

anh

dsp

pre

py 1

po

mt

kao o

o

200 C 100

alb

S

mic

10 hem 1 mt

anh dsp mus

pre

kao

alb mic 10

py

alu

200 C

qtz

100

qtz

po

and

pyr

0.1

alu

S

hem mus

py

anh

dsp

mt 1 pre

kao o

o

300 C

300 C 100 10 mt

po

alb hem

1 0.1

mic

S mus

anh

kao pyr 1

100

dsp

and alu 10

alb dsp

mic

qtz py

pre

1533

10 pre kao mt 1

qtz alu

hem

mus

S anh py

and pyr

1

10

Condensate/Rock (kg/kg) Fig. 5. Amount of precipitated minerals (in grams) as a function of condensate/rock ratio at different temperatures and SO2/H2S ratio in the volcanic vapor. Curve for alunite (alu) is the sum of alunite + natroalunite. A sharp transition between assemblages typical of K metasomatism and advanced argillic alteration can be seen; the position depends on the initial SO2/ H2S ratio in the condensed volcanic vapor. Abbreviations listed in Table 2.

regime at high condensate/rock ratio (e.g., in a fluid-filled fracture) to a rock-dominated regime at low condensate/rock ratio, is shown in Figure 5. The mineral evolution is shown for the two different initial SO2/H2S ratios in the condensed vapor and at three different temperatures: 100°, 200°, and 300°C. At each step, a small amount of fresh rock (we use a 20-g increment, from w/r = 50 to w/r = 0.5) is added to the condensate equilibrated in the previous step of the condensate rock system. The most important feature of this process is the existence of a sharp transition between two different mineral assemblages, an advanced argillic assemblage at high w/r ratio, changing to one in which albite, microcline, and muscovite are stable (Fig. 5). The position of this transition depends on the initial SO2/H2S ratio; the stability of the K metasomatic assemblage occurs over a wider range of w/r for the more reduced fluid (w/r up to ~5), whereas for

the oxidized, SO2-rich fluid this transition occurs at w/r >H2S

-8

o

200 C

-10 o

-12

100 C

Buffering by rock

1

10 Condensate/rock (kg/kg)

Fig. 6. Redox parameter RH (mole fraction) as a function of condensate/ rock (w/r) ratio at different temperatures and initial SO2/H2S ratio. Solid lines correspond to SO2 = H2S (molal); dashed lines are for SO2 >>H2S in the condensed volcanic vapor with total sulfur content the same as where SO2 = H2S. For w/r >5, the redox conditions become governed by the aqueous sulfur system (SO42–- HSO4– - HS– - H2S(aq)). The transition between buffer control by the rock and aqueous solution for this system can correspond to nearly one order of magnitude variation in the condensate/rock (w/r) ratio (with a shift toward the liquid buffer occurring at lower w/r ratio for SO2 >>H2S).

of the mineral assemblage marked by the sharp transition to more reduced conditions. For the oxidized fluid with SO2 >>H2S (dashed lines, Fig. 6), the first irregularity occurs at w/r ~1, where hematite and magnetite disappear sharply (Fig. 5). The second transition

takes place at w/r ~2, where the transition from K metasomatism to advanced argillic alteration occurs. For the reduced fluid (SO2 = H2S), the first transition in RH occurs at w/r ~0.7, where again Fe oxides become unstable, whereas the second transition occurs at w/r ~5, to the equilibrium with advanced argillic alteration. There is one more small transition in RH for the reduced fluid, coinciding with a weak irregularity in the S and pyrite behavior (Fig. 5). Note that there are two ranges of w/r where the term “buffer” makes sense: very low w/r, in which the rock has a buffering capacity (by Fe-bearing minerals), and very high w/r, where the sulfur system of the solution itself buffers the H2 concentration. The rock buffer is stronger with respect to the more reduced fluids, with a sharp transition at higher w/r, ~5. A general picture of the redox state of the modeled system is presented in Figure 7. This now widely used diagram, first proposed by Giggenbach (1987), was originally devised to plot directly analyzed H2 concentrations in active hydrothermal fluids (either in solution or in vapor) versus the measured temperature of the thermal feature (Fig. 2). A set of buffer curves represent different mineral or fluid redox buffers, including the gas and rock buffers (Figs. 2, 7). The range of RH values in the modeled fluids at equilibrium (Fig. 7), noted by the green ruled area, is almost coincident with the magnetite-hematite buffer. Discussion Comparison of modeled species with Satsuma Iwojima acidic springs The condensate of vapor from Satsuma Iwojima fumaroles was reacted in the model with Satsuma Iwojima rhyolite from 350° to 100°C at increasing condensate/rock ratios. Comparison of the ed 100°C solution composition agrees well with the compositions of the acidic hot springs around the shores of

Table 3. Calculated Solutions (100°C) at Various Condensate/Rock Ratios, as Well as Diluted Case of 10:1 W/R, Plus Compositions of Satsuma Iwojima Hot Springs and Kawah Ijen Crater Lake 100°C liquid at condensate/rock

Oxidation state

pH

Cl

SO4

2:1 SO2 = H2S 4.3 12,409 SO2 >>H2S 2.6

2,845 6,795

Al Fe K Na Ca H2S 0.93 56,696 6,652 205 137 5,336 3 10:1 SO2 = H2S 1.39 12,408 1,242 1,184 0.3 81 2,634 145 SO2 >>H2S 0.55 75,872 6,513 301 932 3,062 192 10:1 case: 1.15 3,102 18,968 1,628 75 233 766 48 SO2 >>H2S, diluted 1:3 Satsuma Iwojima springs Kawah Ijen crater lake

1.1−1.7 0.18

347−2,960 20,175

4,390−18,680 396−1,440 50,555

4,550

100−651

95−209

265−1,436

127−315

2,000

1,892

1,685

883

Notes: Aqueous species (ppm) and pH at 100°C after model reaction of vapor condensates with rhyolite (Table 1) at various condensate/rock ratios (Fig. 3b); S° present in all products; two values are listed for the three model results (for 2:1, 5:1, and 10:1 condensate/rock ratio), based on SO2 = H2S and SO2 >>H2S (Fig. 3b), except for Cl, which is not affected by the oxidation state of the S; the 10:1 model results for SO2 >>H2S (values in italics) are diluted 1:3 with ground water and listed for comparison with the range of Satsuma Iwojima spring compositions, which are mixtures of condensate and ground water in ratios of 1:1 to 1:7 (Hedenquist et al., 1994), averaging about 1:3; Kawah Ijen crater lake composition is from Delmelle and Bernard (1994); the high concentrations are partially due to evaporation



MODELING OF HYPOGENE ADVANCED ARGILLIC ALTERATION BY VOLCANIC VAPOR CONDENSATES

1535

Fig. 7. Redox diagram for magmatic-hydrothermal fluids (after Giggenbach, 1987), similar to Figure 2 but with model results plotted. Red star shows the redox state of the high-temperature Satsuma Iwojima volcanic vapor, as sampled from fumaroles at 505° to 877°C (Shinohara et al., 1993). The rock (FeO-FeO1.5) and magnetite-hematite (Mt-Hem) redox buffers are shown by blue and pink lines, respectively, in equilibrium with vapor and liquid phases; the composition of two-phase conditions is represented by pink- and blue-ruled areas. The green-ruled area corresponds to RH values of the modeled solutions, i.e., vapor condensates that cool and react with rhyolite and form alteration minerals. The main tendencies for the modeled RH values are to become more oxidized at higher w/r and SO2/H2S ratios (green arrow). Also shown are the S-gas buffers for the vapor phase and liquid (aqueous solution) phase (Giggenbach, 1987; Einaudi et al., 2003). The modeled solution compositions mainly correspond to (i.e., are largely constrained by) the field of RH values between the rock and Mt-Hem buffers for the liquid phase. See text for more discussion; abbreviations listed in Table 2.

Satsuma Iwojima (Table 3), after correcting the model condensate results for the observed ground-water dilution that averages ~75% (~1:3; Hedenquist et al., 1994). Using the results (in italics) for a w/r ratio of 10:1 and SO2 >>H2S (15:1, the ratio measured in Satsuma Iwojima fumaroles; Table 1), the agreement (Table 3) with the range of Satsuma Iwojima spring compositions supports the overall validity of our model calculations. An exception to this agreement is the low concentration of modeled Fe2+ and Ca2+ (up to factors of 2 for the SO2 >>H2S calculation but much larger for SO2 = H2S). In the case of the Ca discrepancy, the modeled abundance of anhydrite within the residual quartz zone does not agree with its observed paucity in lithocaps, although anhydrite is present in the marginal halo and veins of some ore deposits and is a late-stage, postore mineral that locally fills vugs in residual quartz (Hedenquist et al., 1998; Claveria, 2001). However, massive anhydrite is present beneath the margins of the crater of Kudryavy volcano, Kuril Islands, where hyperacidic condensates of high-temperature (≤940°C) fumaroles infiltrate through the crater floor (Taran et al., 1995; Korzhinsky et al., 1996). Zimbelman et al. (2005) reported the common presence of anhydrite (and/or gypsum) in veinlets accompanying advanced argillic alteration

that is exposed in the eroded portions of several active volcanoes in western North America. In addition, anhydrite is a common alteration mineral with alunite in the El Volcán gold deposit, Maricunga belt, northern Chile (Bernd Lehmann, pers. commun., 2011). Thus, it is likely that anhydrite was more common than is now observed in lithocaps, probably because it has been dissolved during late-stage and postmineral fluid flow (Cooke and Simmons, 2000); such dissolution may be partly responsible at Satsuma Iwojima for the elevated Ca2+ in the acidic springs. Another cause of this disagreement is likely due to the closed-system, single-step fluid-rock interaction that was modeled. For more realistic modeling of actual hot spring compositions, a multistep, flow-through approach is needed (Fouillac et al., 1977; Grichuk, 1988), where a batch of the initial condensate first passes through a column of the primary rock in a temperature gradient, then a second batch flows through the initially altered column, and so on. Realization of such a model should be the next step for quantitative modeling of a specific volcanic hydrothermal system such as Satsuma Iwojima. Acidic solutions (volcanic springs and crater lakes) are considered to be “immature,” as they form through congruent rock

1536

HEDENQUIST AND TARAN

dissolution, aside from the silica (Giggenbach, 1982, 1988). In most cases their cation (metal) ratios are close to that of the unaltered rock matrix, in contrast to mature, equilibrated waters of magmatic-meteoric hydrothermal systems (Giggenbach, 1988), where cation activities are controlled by mineralfluid equilibria. In our model full equilibrium is attained, and cation activities are therefore controlled by the calculated equilibrium. At low pH, the activities of the K and Na species are controlled by the presence of alunite and Na alunite or their solid solution; Ca by anhydrite; Fe by pyrite, and Fe (hydro)oxides; and Al by all alumino-silicates that are stable under the given conditions. The real process of equilibration and replacement of primary minerals with a stable alteration assemblage first passes through dissolution and then crystallization; both processes can be decoupled in time and space. In most cases the acidic waters of acidic crater lakes and hot springs, including the less diluted springs of Satsuma Iwojima, are supersaturated with respect to anhydrite and alunite at any temperature, which means an excess of Na, K, Ca, and Al compared to water-rock equilibrium compositions. Comparison with alteration zonation observed in lithocaps The calculated alteration sequence related to cooling of a condensate of volcanic vapor that is increasingly reactive matches the pattern of upward flaring of alteration zones

observed in lithocaps, the hosts to high sulfidation ore deposits (Steven and Ratté, 1960; Sillitoe, 1999), as well as in their higher temperature roots that are the transition to the porphyry environment (Sillitoe, 2010). Within the high sulfidation environment, ore is typically hosted by the most strongly leached rock, which is dominated by residual quartz with up to >99 wt % SiO2. In structurally controlled high sulfidation deposits, the upward-flared zone of residual quartz, commonly with a remanent vuggy quartz texture due to the remanent texture produced by leached phenocrysts (Steven and Ratté, 1960), is the principal host to subsequent Au mineralization because of its intrinsic permeability. Quartz-alunite alteration is present as a nearly ubiquitous halo to the vuggy quartz core, as well as beneath it, at Summitville and in most similar high sulfidation deposits (Arribas, 1995); at greater depths (and higher paleotemperature), pyrophyllite is common near the feeder zones (e.g., MacDonald et al., 2011). Similar quartzalunite zoning is also noted in flat-lying lithocaps, albeit above and below the silicic horizon that is generated by lateral fluid flow (Fig. 1), in contrast to the alteration halos on the sides of the structural feeder zones (Gonzalez, 1956; Steven and Ratté, 1960; Harvey et al., 1999; Sillitoe, 1999). The upward increase in the widths of the advanced argillic alteration zones along structures (Fig. 8) is caused by cooling of the acidic condensate during ascent and the resulting

dilution quartz – alunite residual decrease in w/r

150 oC pH>2

permeable lithologic horizon (host to lithocap)

quartz 200 oC pH~0.7

250 oC pH~1 300 oC pH~2

350 oC vapor e condensate

+

brine (potassic)

+

500 oC

+

~200 m

+ +

+

+ +

+

+ +

+

Fig. 8. Schematic cross section of advanced argillic alteration due to vapor condensate ascending along a structure; the alteration widens upward due to an increase in reactivity (lower pH) at lower temperature (Fig. 3b; example shown here for cooling from 300°−200°C, where SO2 >>H2S and w/r ~10:1). If the structure intersects a permeable lithologic horizon, lateral flow accompanied by further cooling (dashed isotherms) and increased reactivity (lower pH) will form a lithocap of residual quartz with a halo of quartz-alunite. In distal positions, the w/r ratio decreases, and there is also increased dilution by ground water, both of which cause the reactivity to decrease (Fig. 3b), and alunite to become stable again (changing from w/r ~10:1 to ~2:1; Fig. 4). This is the pattern observed in the quartz-alunite lithocap and residual quartz core at Lepanto (Fig. 1). Thus, the main body of advanced argillic alteration near the surface may be on the shoulder of the porphyry deposit, offset from the deep intrusive source of acidic volatiles (Fig. 1). Potassic alteration associated with the porphyry forms at the same time as the shallow advanced argillic alteration, due to the coupled hypersaline liquid (brine) and vapor, respectively (Hedenquist et al., 1998).



MODELING OF HYPOGENE ADVANCED ARGILLIC ALTERATION BY VOLCANIC VAPOR CONDENSATES

increase in reactivity (Fig. 3b). Alunite typically formed at ~220° to 280°C in high sulfidation deposits, based on an assessment of S isotope systematics for alunite-sulfide pairs (Arribas, 1995); 250°C is equivalent to a paleodepth of ~400 m in a boiling hydrostatic system, deeper in nonboiling systems. This modeling study shows that alunite is dissolved at temperatures of ~