13.

15 Geochemistry of Hydrothermal Gold Deposits

JA Saunders, Auburn University, Auburn, AL, USA
AH Hofstra, U.S. Geological Survey, Denver, CO, USA
RJ Goldfarb, U.S. Geological Survey, Denver, CO, USA
MH Reed, University of Oregon Eugene, OR, USA
ã 2014 Elsevier Ltd. All rights reserved.

13.15.1 Introduction 383

13.15.2 Epithermal Deposits 384
13.15.2.1 Introduction 384
13.15.2.2 Low-Sulfidation Epithermal Deposits 385
13.15.2.2.1 Trace elements and mineral associations 385
13.15.2.2.2 Ore fluid characteristics 386
13.15.2.2.3 Hydrothermal alteration 386
13.15.2.2.4 Geochemistry of the Midas LS deposit, Nevada 387
13.15.2.3 High-Sulfidation Epithermal Ores 390
13.15.2.3.1 Hydrothermal alteration 390
13.15.2.3.2 Trace elements and mineral associations 391
13.15.2.3.3 Ore-fluid composition 392
13.15.2.3.4 Light-stable isotopes 393
13.15.2.3.5 Reaction of volcanic gas condensate with quartz latite 393
13.15.2.3.6 Ore deposition 394
13.15.2.4 New Ideas about the Geochemistry of Epithermal Deposits 397
13.15.3 Carlin-Type Gold Deposits 398
13.15.3.1 Introduction 398
13.15.3.2 Age and Geologic Setting of CTD 399
13.15.3.3 Ore and Gangue Minerals 399
13.15.3.4 Geochemistry of Rocks and Pyrite 400
13.15.3.5 Composition of Ore Fluids 404
13.15.3.6 Carbonate Dissolution 404
13.15.3.7 Ore Fluid Composition and Precipitation Mechanisms 405
13.15.3.8 Element Substitution in Pyrite 406
13.15.3.9 Source(s) of Ore Fluid Components 407
13.15.4 Orogenic Gold Deposits 409
13.15.4.1 Introduction 409
13.15.4.1.1 General geologic setting and genetic model 409
13.15.4.2 Geochemistry and Mineralogy of Alteration and Ores 411
13.15.4.2.1 Ore mineral assemblages 411
13.15.4.2.2 Wall rock alteration 412
13.15.4.2.3 Geochemistry of ore-forming fluids 413
13.15.4.2.4 PTX constraints on ore deposition 415
13.15.4.3 Geochemistry of Type Examples 416
13.15.4.3.1 Phanerozoic metasedimentary-hosted Muruntau 416
13.15.4.3.2 Paleoproterozoic BIF-hosted Homestake 417
13.15.4.3.3 Late Archean Greenstone-hosted Golden Mile 418
13.15.5 Summary and Conclusions 418
Acknowledgments 419
References 419

13.15.1 Introduction geotectonic settings and relationship to other important de-

posit types are summarized in Figure 1.
The geochemistry of the three most significant hydrothermal Another important hydrothermal gold deposit type, por-
gold deposit types, epithermal, Carlin-type, and orogenic, is phyry copper–gold deposits, is not discussed here. Among
described in this article. The geologic characteristics of these the three deposit classes considered, the high-sulfidation
deposits are discussed in detail elsewhere (Cline et al., 2005; epithermal and probably most low-sulfidation epithermal de-
Goldfarb et al., 2005; Simmons et al., 2005), and their general posits are genetically linked to coeval magmas and magmatic

Treatise on Geochemistry 2nd Edition http://dx.doi.org/10.1016/B978-0-08-095975-7.01117-7 383

384 Geochemistry of Hydrothermal Gold Deposits

Continent Oceanic Back Accreted Continental Back-arc

ridge arc terranes arc extension

Epithermal VHMS Orogenic Epithermal Epithermal/

Au Cu–Au Au Au hotspring
Porphyry Porphyry Au
Cu–Au Au–Au Carlin-style
(±skarns) (±skarns) Au

Accretionary wedge Continental crust Granitoids

Oceanic crust Subcrustal lithosphere Asthenosphere

Compressional fault thrust Extensional fault Basalt

Figure 1 Diagrammatic illustration showing the geotectonic settings of the gold deposits discussed here and some associated ores types (Groves
et al., 1998).

fluids, although it is clear that there is a large component of igneous rocks at temperatures and pressures that he inferred to
meteoric water in the ore-forming solutions (Saunders et al., be less than 200  C and
100 bars. The term is still applied to
2008). Carlin-type deposits (CTD) might also have a magmatic ore settings where a link to a magmatic fluid source is implied
fluid source although the magmatic connection is still seriously by the proximity of the ores to contemporaneous volcanic or
debated (Cline et al., 2005; Emsbo et al., 2006). Magmatic intrusive rocks, but, in practical terms, epithermal ores are
fluids are not involved in the third deposit class considered distinguished by particular assemblages of ore minerals and
here, orogenic gold, which formed from metamorphic fluids wall rock alteration minerals and by ore textures and structures
released along active collisional or compressive plate bound- that indicate a shallow crustal setting. They are most com-
aries (Goldfarb et al., 2005; Groves et al., 1998). Together, monly mined for gold and silver, although copper, lead, and
deposits of these three ore types account for
75% of the zinc are also abundant in many such ores. Modern fluid inclu-
gold mined globally from lodes. sion microthermometry (Hedenquist and Henley, 1985; Nash,
Since the time of Boyle’s (1979) description of the geo- 1972; Simmons et al., 2005) has extended the temperature
chemistry of the entire spectrum of gold deposits, there have range of epithermal ores to about 300  C and indicates typical
been some significant advances in knowledge based in part on pressures up to many hundreds of bars.
new geochemical analytical techniques. These include multi- Epithermal precious-metal ores occur widely in Tertiary
element inductively coupled plasma optical emission spec- and younger rocks throughout the Pacific Rim and in other
trometry (ICP-OES) and ICP mass spectrometry (ICP-MS) Phanerozoic orogenic belts worldwide. The shallow setting of
analysis; high-precision and in situ stable and radiogenic iso- epithermal ores leads to their relatively rapid erosion and
topic analyses; development of stable isotopic techniques for skews the ages of known deposits to Cretaceous and younger,
zinc, copper, selenium, and iron; new geochronologic tools but many older deposits are known. They are tied to volcanism
such as single-crystal zircon U–Pb and 40Ar/39Ar analyses; in a variety of geotectonic settings and exhibit a wide range of
and laser ablation (LA)-ICP-MS of minerals, and single fluid mineralogical and geochemical characteristics. Epithermal ores
inclusions to measure dissolved species. The new analytical have been classified a number of different ways but most
techniques have led to a better understanding of the geochem- recognize that there are two end-member deposit types: (1)
istry not only of the ores and fluids but also of the ore-forming deposits that occur above coeval magmas, typified by advanced
processes, which can now be well constrained by computer argillic alteration with gold, enargite, covellite and pyrite,
modeling founded on a wealth of new thermodynamic data. and fluid inclusions of moderate to high salinity; and (2) de-
posits in veins, breccias, and stockworks that typically flank
a magmatic center, where gangue minerals such as adular-
13.15.2 Epithermal Deposits ia, sericite, and carbonates occur with ores containing elec-
trum, acanthite, and silver-bearing sulfosalts, as well as
13.15.2.1 Introduction
locally abundant selenide and telluride minerals, all with
The term epithermal was coined by Lindgren (1933) to refer to fluid inclusions of low salinity. Between these end-members
deposits formed near the surface in genetic connection with are epithermal vein ores commonly dominated by sulfides of
 Geochemistry of Hydrothermal Gold Deposits 385

zinc, lead, copper, and silver with additional silver and gold in Nevada Rift (NNR) but also proposed that the NNR was formed
electrum, and fluid inclusions of moderate salinity. In the by the Yellowstone hotspot.
terminology of Hedenquist et al. (2000) and Sillitoe and There are important exceptions to the general rule that LS
Hedenquist (2003), the first category mentioned earlier is ores are formed during regional extension. For example, the
termed high-sulfidation (HS) and the second is called low- Plio-Pleistocene McLaughlin deposit formed in a transpressive
sulfidation (LS), and the transitional ore type between these setting within the San Andreas fault system (Sherlock et al.,
two is termed intermediate sulfidation (IS). In this article, 1995; Sillitoe and Hedenquist, 2003). In summary, LS ore
focused on gold ores, the geochemistry of the two end-member formation may have deep roots, apparently initiated by partial
types is described because gold is the economically dominant melting of a ‘prepared’ lithospheric mantle (Richards, 2009;
precious metal in these two forms of mineralization. Although Saunders and Brueseke, 2012).
the authors apply the ‘sulfidation’ classification terms here
because they have been widely adopted, alternative classi-
fications that distinguish the same categories using descrip- 13.15.2.2.1 Trace elements and mineral associations
tive schemes are also widely used, as outlined by Simmons The LS gold deposits are commonly enriched in Ag, As, Au,
et al. (2005). Hg, Sb, Se, Te, and locally Tl, a group of elements generally
known as the ‘epithermal suite.’ Base metals, Cu, Pb, and Zn,
are also locally abundant and some ores become relatively
13.15.2.2 Low-Sulfidation Epithermal Deposits
enriched with the base metals at deeper levels in veins and
The global abundance of LS ores may reflect the numerous thus resemble the IS ore type. In addition, Mo, Sn, V, and W
geotectonic settings where they form. Most LS ores formed in may be enriched in LS ores and in their altered wall rock
settings where earlier subduction enriched the lithospheric (Jensen and Barton, 2000; John et al., 2003; Simmons et al.,
mantle above the subducted plate with volatiles and metals. 2005). In the northern Great Basin of the USA, chalcopyrite is
Partial melting of this subduction-enriched mantle wedge pro- particularly abundant in the highest grade LS ores (Saunders
duced magmas that yielded porphyry and epithermal mineral- et al., 2011a). This is true also in the bonanza LS gold ores
ization (e.g., Richards, 2009; Saunders and Brueseke, 2012). at Hishikari (Figure 2; Izawa et al., 1990; Nagayama, 1993;
The LS ores are typically hosted by basalt–rhyolite calc-alkaline Shikazono and Nagayama, 1993), Kushikino (Matsuhisa et al.,
volcanic rocks in predominantly extensional environments in 1985), and Koryu (Shimizu et al., 1998a, 1998b). At Hishikari,
intra-arc, near-arc, and back-arc postcollisional rifts (Sillitoe there is a stronger statistical correlation between Au and Cu than
and Hedenquist, 2003). The LS ores are generally similar in there is between Au and Ag (Nagayama, 1993). HS ores are also
age to the nearby or host volcanic rocks (Hames et al., 2009; very much enriched in Cu, as discussed in the succeeding text.
Izawa et al., 1990; Leavitt et al., 2004; Tohma et al., 2010). New Because of the abundance of telluride and selenide minerals in
high-precision geochronology suggests that ore formation was some LS ores, Lindgren (1933) separated them into separate
of relatively short duration, perhaps as brief as a few thousand epithermal subclasses. Typically tellurides and selenides do not
to a few tens of thousands of years (Heinrich, 2006; Saunders, occur together, but there are exceptions such as at the Prasolovs-
2010; Simmons and Brown, 2006; Tohma et al., 2010). Some koye deposit, Kunashir Island, Kuril island arc (So et al., 1995).
large LS deposits (Ladolam, Porgera, and Cripple Creek) In the telluride ores of Colorado, calaverite (AuTe2), sylvanite
formed in association with alkaline rocks in a variety of (AuAgTe4), krennerite [(Au,Ag)Te2], petzite (Ag3AuTe2), hessite
convergent-margin extensional settings, as opposed to well- (Ag2Te), native gold or electrum, and tetrahedrite are the main
developed rifts, and have some geologic characteristics some- ore minerals, but native Te and tellurides of Fe, Ni, Pb, and Hg
what atypical of most LS deposits (Richards, 1995; Sillitoe and are also locally abundant (Saunders, 1988, 1991; Saunders and
Hedenquist, 2003). May, 1986). An interesting geochemical and mineralogical as-
The diversity of geotectonic and magmatic processes respon- sociation in Colorado, as well as worldwide for LS telluride-rich
sible for forming LS ores in a single setting is well illustrated by ores, is the ubiquitous presence of the V-bearing mica roscoelite.
the Miocene–Pliocene deposits of the northern Great Basin of In a study of several Colorado LS telluride ores, Saunders (1986)
the USA. John (2001) showed that these LS ores are contempo- showed that the ores were very similar irrespective of whether
raneous with two different host petrogenetic magma suites that they formed during Laramide subduction or later rifting.
reflect variations in the geotectonic setting: (1) continental rift- Saunders (1986) also showed that the country rocks that hosted
related bimodal rhyolite and basalt that is K-rich tholeiite and the ores were highly variable in their composition and age and
(2) a subduction-related andesite > dacite, oxidized high-K, thus apparently were not the source of the precious metals,
calc-alkaline volcanic assemblage. The LS Sleeper, National, Te, or V. However, Jensen and Barton (2000) recognized a
Midas, and Mule Canyon deposits are hosted by the bimodal correlation between the amount of mafic wall rock and V
suite, whereas LS deposits such as the Comstock Lode and content of veins and alteration minerals. Saunders et al.
Tonopah and HS ores such as Goldfield and Paradise Peak are (2008) concluded that the host rocks were not the source of
in the subduction-related andesite assemblage (John, 2001). Au, Ag, and Se in the bonanza LS ores of the northern Great
Saunders et al. (1996, 2008) and Kamenov et al. (2007) pro- Basin; the precious metals and Se apparently are derived
posed that the 16–14 Ma LS ores in northern Nevada and from deeper mafic magmatic sources based on Pb isotope data
southwestern Idaho associated with the bimodal volcanic (Kamenov et al., 2007). However, country rocks in the upper
assemblage were genetically related to the initial emergence of crust apparently did contribute some major chemical con-
the Yellowstone hotspot mantle plume. John (2001) concluded stituents in the ores (e.g., silica, Na, K, Ca, and Al; Saunders
that these LS ores were more directly related to the Northern et al., 2008).
386 Geochemistry of Hydrothermal Gold Deposits

Figure 2 (Left) Photograph of banded Ryosen no. 6 vein, Hishikari deposit, Japan, which is up to 2 m thick. (Right) Photograph of the Colorado Grande
vein at the Midas deposit, Nevada, which is approximately 1 m thick in the image. Note fragments of wall rocks encrusted by silica in the vein.

13.15.2.2.2 Ore fluid characteristics 13.15.2.2.3 Hydrothermal alteration

Studies of fluid inclusions trapped in gangue minerals show Hydrothermal alteration in LS deposits can be subdivided into
that LS ores, or at least the gangue minerals, typically formed at (1) K-feldspar or sericite- and chlorite-altered wall rock border-
150–300  C, from dilute solutions, typically 0.1 to
5.0% ing veins at depth; and (2) the overlying draping blanket of
NaCl equivalent (Hedenquist and Henley, 1985; Nash, 1972; pervasive intermediate and advanced argillic alteration and
Simmons et al., 2005). The more telluride-rich LS examples silicification. The latter is produced by downward percolating
typically have slightly higher dissolved salt contents of 5–9% near-surface waters acidified by oxidized H2S gases and by CO2
NaCl equivalent (Ahmad et al., 1987; Saunders, 1988, 1991; condensation, both accompanied by cooling of silica-rich wa-
Saunders and May, 1986). However, telluride-rich ore-forming ters. At the district scale, there is additional regional propylitic
fluids are often CO2 rich, which depresses the melting point of alteration (not specifically related to precious-metal minerali-
ice in the inclusions such that the salinity value will be errone- zation) consisting of chlorite, epidote, and calcite that is par-
ously high unless the CO2-clathrate melting temperature is ticularly evident in volcanic host rocks.
used to determine the salinity (Hedenquist and Henley, Hydrothermal alteration associated with LS ores is inter-
1985). The fluid inclusion observations often show that the preted as a product of complex water–rock–vapor interactions.
ore-forming solutions were boiling, which is an effective mech- For example, acid volatiles released from a cooling magma
anism for depositing precious metals from solution (Spycher chamber or boiling solutions at depth are incorporated into
and Reed, 1989) and is also commonly observed in geother- the heated groundwater comprising the geothermal system
mal systems at depth (e.g., Hedenquist, 1991). The low salin- where an epithermal vein forms. This decreases the pH of the
ities of fluid inclusions are largely consistent with H–O isotope geothermal waters and can lead to local acid argillic alteration
ratio measurements, which show that the vein-forming solu- or even complete destruction of the host rocks at very low pH,
tions are predominantly heated groundwaters with some de- such as in the mud pots of advanced argillic alteration at
gree of oxygen isotopic exchange with the surrounding country Yellowstone National Park. This local advanced argillic alter-
rocks, as occurs in active geothermal systems (Hedenquist, ation, which forms the common steam-heated blankets in the
1991; Simmons et al., 2005). In fact, at the <1 Ma Hishikari LS epithermal environment, should not be confused with the
deposit (Figure 2), where veins formed at
200–220  C (Etoh pervasive advanced argillic alteration intimately associated
et al., 2002; Izawa et al., 1990), 70  C water is still discharging with HS ores discussed in the succeeding text. In LS ores,
along the vein structure. In contrast to the predominant advanced argillic alteration is caused by oxidation of H2S
meteoric-water source in the ore-forming fluids, detailed sulfur at and above the water table to H2SO4 (Hedenquist, 1991).
isotope studies (Jensen, 2003; John et al., 2003; Shelton et al. Hydrothermal alteration of the country rock may be produced
1990; So and Yun, 1996; Yuningsih et al., 2011; Zhang and prior to ore deposition by lower-pH solutions than those that
Spry, 1994) on LS epithermal ores suggest that sulfur in sulfide deposit the precious metals. For example, there is widespread
minerals associated with precious-metal minerals has a mag- pervasive alteration of the volcanic rocks at Cripple Creek,
matic source (e.g., d34S values in the 5 to þ5% range) or Colorado (Jensen and Barton, 2000), but lamprophyre breccia
perhaps a mixture of magmatic and host-rock sulfur (Castor fragments coated with ore and gangue minerals, such as cala-
et al., 2003). verite, petzite, dolomite, and quartz, deep in the Cresson
 Geochemistry of Hydrothermal Gold Deposits 387

diatreme at Cripple Creek (2200 level in the 1930s mining), deprotonated equivalent, bisulfide (HS). The H2S(aq) is the
are virtually unaltered, suggesting that ore-forming solutions quintessential ‘soft Lewis acid’ and, consequently, makes
were not very corrosive at a relatively low temperature of strong covalent bonds with soft Lewis bases, such as Au, Ag,

185  C (Saunders, 1988). The Bessie G Au–Ag–Te deposit Hg, As, Sb, Te, and Se (Langmuir, 1997; Saunders et al., 2008).
in Colorado is another example of very weak hydrothermal The implications of this are threefold: (1) H2S-rich fluids will
alteration associated with LS ore-forming solutions. At Bessie dissolve significant amounts of Au and Ag, although how
G, the host rocks for the vein are Mesozoic siltstones and much is debated (see Brown, 1986; Saunders, 2010; Saunders
sandstones, and associated silicification extends only a few and Schoenly, 1995; Simmons and Brown, 2006); (2) H2S in
centimeters into the wall rocks (Saunders and May, 1986). It the ore-forming fluid causes the epithermal suite of elements to
is likely that hydrothermal alteration of wall rocks provides travel together; and (3) high levels of dissolved H2S suppresses
most of the chemical constituents to form the non-ore minerals the solubility of base metals such as Fe, Pb, Zn, and Cu
in the veins (Saunders et al., 2008). For example, at the Sleeper (Saunders and Schoenly, 1995). A number of physicochemical
deposit in Nevada, glassy rhyolites hosting the ore were the processes can lead to Au precipitation from a neutral pH,
source of high concentration of silica in ore-forming fluids that moderately reduced, H2S-rich ore-forming fluid, including
deposited opal (Figure 3) in the high-grade veins (Saunders and boiling, mixing with cooler and more dilute groundwaters,
Schoenly, 1995). In contrast, the Mule Canyon deposit in reaction with wall rocks, particularly iron minerals, oxidation,
Nevada, hosted predominantly by basalts, lacks significant and simple cooling. All of these processes may be important
amounts of silica in the ores (John et al., 2003). Saunders et al. locally in LS ores and, in some cases, more than one may
(2008) observed that some very adularia-rich LS epithermal veins induce ore mineral precipitation. Boiling (Drummond and
in Nevada, such as those at Jumbo, Tenmile, and Sandman, were Ohmoto, 1985; Romberger, 1988; Saunders and Schoenly,
hosted by metapelites and proposed that these were the source 1995; Spycher and Reed, 1989) can be the paramount process
for the Al, Si, and K. Leaching of Al from wall rocks requires low- in LS ore formation, particularly when the solutions are signif-
pH geothermal waters presumably caused by absorption of acid icantly undersaturated with respect to gold (Brown, 1986;
volatiles from below. However, most LS ores appear to have Spycher and Reed, 1989).
formed from near-neutral pH solutions. This apparent disparity Taylor (2007) posed the question as to whether Au-rich
can best be explained by widespread boiling in fractures that epithermal deposits form mainly due to gold-enriched sources,
leads to a significant pH increase (Drummond and Ohmoto, such as fluids or magmas, or from exceptionally efficient mech-
1985; Spycher and Reed, 1989) primarily because of loss of CO2. anisms or processes of Au precipitation. Saunders (2010) ar-
Thus, boiling causes platy calcite formation (Simmons and gued for the latter possibility, at least in the formation of
Christensen, 1994), precious-metal precipitation by H2S loss, bonanza epithermal ores, based on new data indicating how
base metal precipitation by pH increase, and deposition of silica, short-lived epithermal ore-forming processes can be. Neverthe-
adularia, and sericite. less, in support of the former, Saunders and Schoenly (1995)
The general geochemistry of LS ore-forming fluids is fairly found that ore-forming solutions at the Sleeper deposit may
well established (see Reed, 1997; Saunders and Schoenly, have contained as much as several ppm of dissolved gold,
1995; Spycher and Reed, 1989). Ore-forming solutions are which was about three orders of magnitude higher than ob-
typically near-neutral pH, consistent with deposition of adu- served in the geothermal solution from Ohaaki-Broadlands,
laria, sericite, and/or carbonate, moderately reducing (pyrite- New Zealand (Brown, 1986). If one assumes a boiling geother-
stable), and contain significant amounts of H2S and minor mal system is generally part of an LS epithermal ore-forming
CO2. Gold and silver in these ore-forming fluids typically system (but not necessarily directly related to ore deposition;
form stable aqueous complexes with H2S(aq) and its e.g., Saunders et al., 2008), then what additional requirement
is needed to make a bonanza epithermal deposit and not just a
low-grade or barren geothermal system? Perhaps the answer
lies in understanding controls on the formation of gold- and
volatile-rich, shallow magmas that episodically release their
metalloids and volatiles within the porphyry environment.
Increasingly, there appears to be a connection between hydro-
thermal processes in the deeper porphyry setting and much
shallower epithermal setting (e.g., Hedenquist and Lowenstern,
1994; Heinrich, 2005; Heinrich et al., 2004).

13.15.2.2.4 Geochemistry of the Midas LS deposit, Nevada

The Midas deposit of northern Nevada consists of LS epither-
mal veins that are chosen here as a representative example of LS
0.5 mm epithermal systems because of its excellent documentation and
its characteristics that exemplify LS ores. Details of the general
Figure 3 Photograph of an opal-electrum band encrusted by later geology the Midas district and veins are summarized here from
silica-only bands from the Sleeper deposit, Nevada. In detail, the bands Goldstrand and Schmidt (2000), Leavitt et al. (2004), Leavitt
consist of three-dimensional dendrites of electrum formed by and Arehart (2005), Riederer (2007), and Chitwood (2012).
aggregation of colloids; top of the image is toward the vein center Further, the Midas deposit has experienced little post-ore fault-
(see Saunders, 1990, 1994). Scale bar ¼ 0.5 cm ing or erosion. It is also considered a world-class deposit,
388 Geochemistry of Hydrothermal Gold Deposits

containing a resource of
125 t Au and 1250 t of Ag (Leavitt mineralized rock has an age of 15.2 Ma. The similarity in age
et al., 2004). Although there had been some gold production of the rhyolitic rocks at Midas to the ores led Leavitt et al. (2004)
in the district in the early 1900s (
3 t), it was not until the to propose that rhyolites contributed both heat and metals to
early 1990s that the large veins were discovered at Midas. Not drive the ore-forming solutions. However, Pb isotope data pre-
long after, Saunders et al. (1996) and John (2001) proposed sented by Saunders et al. (2011c) suggests that Pb in electrum at
that Midas was just one of a clan of mid-Miocene bonanza Midas was derived from a mafic rock source, perhaps a deeper
epithermal deposits in the northern Great Basin, including magma chamber that was the source of the mafic sills, dikes,
National, Sleeper, Mule Canyon, Buckskin National, DeLamar, and flows exposed in underground workings in the district.
and Ivanhoe that formed during the volcanism attending the Major veins at Midas include the Colorado Grande and
initial emergence of the Yellowstone hotspot. Discovery, along with the Gold Crown vein that is a footwall
From a geochemical standpoint, Midas and similar deposits split off the Colorado Grande vein. Recent drilling and under-
of the northern Great Basin contrast to some extent with LS ores ground workings have encountered additional, much more
worldwide because of their Se enrichment. Lindgren (1933) silver-rich and smaller veins to the east of the main producing
recognized the importance of these Se-rich ores and proposed veins (Chitwood, 2012). Most of the production from Midas to
they were a subclass of epithermal ores. Although the mid- date has come from the Colorado Grande vein, which has a
Miocene northern Great Basin LS ores are enriched in Se, this strike length of
1.5 km and is typically 1–2 m wide (Figure 2).
does not relate specifically to the Yellowstone hotspot and/or A longitudinal section through the Colorado Grande vein shows
NNR. Saunders and Brueseke (2012) pointed out that similar that mineralization is confined to a vertical interval of
600 m
Se-rich bonanza LS ores occur in the Eocene Republic district of (Figure 4). This elevation control is typical in epithermal de-
Washington and the Plio-Pleistocene epithermal ores of Japan. posits and is seen elsewhere in northern Great Basin, such as at
Thus, Se-rich LS epithermal ores do not necessarily relate to a the Sleeper deposit, and has been attributed to the boiling level
mantle plume, but perhaps require rifting-induced partial melt- (Nash et al., 1995; Saunders, 1994).
ing of a fertile lithospheric mantle wedge (Saunders and The bottom of the zone is controlled by initiation of boiling
Brueseke, 2012) that is enriched in Se. Alternatively, John et al. due to the drop in pressure as ore-forming solutions move
(2003) suggested that black shales could be the source of Se at from deeper levels, whereas the top represents the elevation
the Mule Canyon LS deposit of northern Nevada. where ore-forming solutions were depleted. There is some
Midas exhibits many geologic characteristics common to LS evidence that the spent ore-forming solutions discharged at
bonanza veins, such as (1) well-defined vein structures typi- the surface, as some siliceous sinter is preserved in the Midas
cally exhibiting fine to coarse banding caused by successive ore district (Leavitt et al., 2004) and in the Ivanhoe district 15 km
and gangue mineral deposition on vein walls (Figure 2) indi- to the southeast (Wallace, 2003).
cating multiple episodes of precious-metal deposition; (2) The Midas veins exhibit typical banded black and white
electrum as the principal ore mineral; (3) general lack of sul- encrustation textures closely resembling those at the Hishikari
fides coprecipitated with precious metals during bonanza ore and many other epithermal deposits (Figures 2 and 5). At
formation, although chalcopyrite and pyrite are common lo- Midas, black-colored bands (Figure 5) host most of the ore
cally; (4) local ore mineral textures consistent with colloid minerals as they do in the ‘ginguro’ (silver-black) ore at
aggregation and physical transport of colloids in the ore- Hishikari and other Japanese LS deposits (e.g., Izawa et al.,
forming solutions (e.g., Saunders et al., 2011b); (5) fluid in- 1990). The principal gangue minerals in the Midas veins are
clusion and other evidence (e.g., pyrite replacing marcasite, quartz, chalcedony, adularia, calcite, and fluorite (Goldstrand
formation of bladed calcite that typically then gets replaced and Schmidt, 2000; Leavitt and Arehart, 2005), and calcite has
by silica, adularia precipitation, and elevation controls on ore the bladed appearance common to boiling hydrothermal sys-
zones) that hydrothermal solutions boiled at the time of ore tems (Simmons and Christensen, 1994) and is only locally
deposition; and (6) ore veins and host volcanic rocks are close replaced by silica. The most significant ore minerals identified
in age. The Midas district lies along the eastern edge of the at Midas veins are naumannite (Ag2Se), aguilarite (Ag4SeS),
NNR, but also it and the nearby mid-Miocene Ivanhoe LS chalcopyrite, electrum (Figure 6), and locally fischesserite
epithermal district lie along the northwestern extension of (Ag3AuSe2). Clausthalite (PbSe) and rare Cu– and Cu–Fe sele-
the Carlin Trend that contains numerous Eocene world-class, nides have also been identified, along with more common
sediment-hosted gold deposits described in the succeeding text galena, sphalerite, and pyrite (Goldstrand and Schmidt, 2000).
(John, 2001; Leavitt et al., 2004). Within the individual ginguro bands, naumannite at Midas
The bonanza veins at Midas are hosted by an assemblage of commonly exhibits self-organized dendritic textures similar to
mid-Miocene bimodal volcanic and sedimentary rocks, includ- those at the Sleeper deposit (Saunders, 1994), which along
ing mafic dikes, sills, and lava flows; rhyolitic tuffs, including with local textural evidence from the naumannite particles
air-fall, ash-flow, and vitrophyres, and flows; and carbona- (Goldstrand and Schmidt, 2000; Saunders et al., 2011b), indi-
ceous and locally tuffaceous siltstone (Goldstrand and cate that naumannite was physically transported and deposited
Schmidt, 2000; Leavitt et al., 2004). The 40Ar/39Ar geochronol- as colloids or nanoparticles. This is also observed at deposits in
ogy by Leavitt et al. (2004) and Hames et al. (2010) has shown the nearby Ivanhoe, Jarbidge, National, and Seven Troughs
that volcanism at Midas initiated with eruption of apparently districts (Saunders et al., 1996, 2011b). Aggregates of nauman-
pre-ore basalts from 16.32 þ 0.17 to 15.51 þ 0.07 Ma, which nite colloids recrystallize, anneal, and grow into coarser-grained
were overlain by rhyolitic flows and tuffs (
15.6 Ma). Bo- minerals, much as aggregates of silica and electrum colloids can
nanza veins formed at
15.4 Ma at Midas and the nearby (Saunders 1990, 1994). Thus, what is commonly observed are
Ivanhoe district. A post-ore rhyolitic tuff overlying the relatively coarse-grained phases such as those in Figure 6.
 Geochemistry of Hydrothermal Gold Deposits 389

Section 2650 N
S N

1850 m 1850 m
Top

1700 m 1700 m
Tbg
Top
1550 m 1550 m

1400 m 1400 m

>10
10–50 1250 m
1250 m 200m
50–100
Drill hole pierce point >100
Goldstrand and Schmidt (2000)

Figure 4 Longitudinal section through the Colorado Grande vein at Midas showing g t1 Au  vein thickness (m) (reproduced from Goldstrand PM
and Schmidt KW (2000) Geology, mineralization, and ore controls at the Ken Snyder Gold–Silver Mine, Elko County, Nevada: In: John DA and
Wallace AR (eds.) Geology and Ore Deposits 2000: The Great Basin and Beyond, pp. 265–287. Reno, NV: Geological Society of Nevada).

el nm

cp

0.2 mm

Figure 6 Photomicrograph (reflected light) of the principal ore minerals

in the Colorado Grande vein: chalcopyrite (cp), naumannite (nm), and
electrum (el).

Figure 5 Photograph of the black ‘ginguro’ bands of the Discovery vein were typically in the 40:60–60:40 range. Geochemically, trace
at the Midas deposit, Nevada. Black ginguro bands contain mostly elements in veins, for the most part, mirror their principal
dendritic naumannite, electrum, and minor sulfide minerals. Vein width mineral constituents. For example, Redak (2005) studied the
shown here is 0.5 m. geochemistry of the Colorado Grande vein, plotted longitudi-
nal sections of the distribution of Au, Ag, Se, As, Cu, and Hg
Electrum and coexisting naumannite–aguilarite account for in the vein, and showed that they all tended to show the
the bulk of the precious metals in the Midas veins currently same elevation control as gold in the longitudinal section of
being mined, and they are closely associated with chalcopyrite Goldstrand and Schmidt (2000; Figure 4). Redak (2005) con-
(Figure 6), which is similar to the National and Ivanhoe dis- cluded that gold was strongly correlated with Ag, Se, and Hg.
tricts, the Silver City district, Idaho (Saunders et al., 2011a), Earlier statistical analyses of the Colorado Grande samples by
and the Japanese LS ores. Silver-to-gold ratios are variable Goldstrand and Schmidt (2000) showed a strong positive
through the main veins but typically range from 10 to 15:1 correlation between Au, Ag, Se, and Cu, which is similar to
(Saunders, unpublished data). Semiquantitative electrum ana- what was observed at Hishikari (Nagayama, 1993). Leavitt and
lyses by Riederer (2007) showed that Au:Ag ratios of the alloy Arehart (2005) evaluated the geochemical trends of major,
390 Geochemistry of Hydrothermal Gold Deposits

minor, and trace elements in the Colorado Grande and Gold Potassic alteration resulted in gains in K, S, Se, P, Ag, As, Sb,
Crown veins, but did not include gold in their statistics. In Hg, and W and losses in Ca, Mg, and Cr in mafic rocks. In
general, geochemical trends reported by Leavitt and Arehart rhyolites, potassic alteration led to an increase in S, Ba, Mg,
(2005) are similar to the earlier studies, but they did note a Mn, P, Se, Ag, Sb, Cu, W, and Bi and losses in Na, Fe, and Zn.
slight increase in Hg and As at the top of the veins. Leavitt and Finally, potassic alteration of felsic volcaniclastic rocks
Arehart (2005) found no increase in base metals with depth in increased S, Al, Ti, Mg, P, Sr, Cr, Cu, and W and depleted
the veins, which commonly occurs in some LS and IS systems Fe, Na, and Ca. The formation of the banded bonanza veins
(Simmons et al., 2005). appears to have coincided with potassic alteration (Leavitt and
Fluid inclusion studies, most recently by Riederer (2007), Arehart, 2005; Leavitt et al., 2004).
indicate that the ore-forming fluids at Midas were in the
range of 200–260  C and were relatively dilute (<1.7% NaCl
13.15.2.3 High-Sulfidation Epithermal Ores
equiv.), which is similar to many other epithermal systems
(Hedenquist and Henley, 1985). The dilute fluids suggest The gold–copper deposits of Summitville (CO, USA) and of
meteoric-water dominance, which is consistent with interpre- the Nansatsu District (southwestern Kyushu, Japan) are typical
tations from light-stable isotope data elsewhere in northern examples of the acid-sulfate covellite–enargite–gold ore type
Nevada, such as Buckskin National (Vikre, 1985), Mule Can- described in early studies by Steven and Ratte (1960), Saito and
yon (John et al. 2003), and Sleeper (Saunders et al., 2008), that Sato (1978), Stoffregen (1987), and Izawa and Cunningham
indicated most of the ore-forming solutions consisted of (1989). The ores have been called Nansatsu-type (Urashima,
exchanged heated meteoric water. Because ore-forming fluids 1975), Nansatsu-Summitville-type, acid-sulfate Cu–Au, and
at Midas were near-neutral pH, based on the presence of calcite HS. The Nansatsu and Summitville deposits are two of a large
and adularia gangue minerals and moderately reducing number of HS epithermal deposits (see Hedenquist et al.,
(pyrite-stable), they would likely have transported dissolved 1994), which include the historic deposit at Goldfield, Nevada
precious metals as aqueous complexes with H2S(aq) and (Ransome, 1909; Vikre, 1989), and the recently developed
bisulfide (HS; e.g., Seward, 1973; Spycher and Reed, 1989). deposits at Pascua, Chile (Chouinard et al., 2005a, 2005b),
Boiling of such a fluid is very effective in precipitating precious and Yanacocha, Peru (Teal and Benavides, 2010). All are
metals (Brown, 1986; Spycher and Reed, 1989). hosted by volcanic rocks and shallow subvolcanic intrusions
As discussed in the preceding text, hydrothermal alteration that are intensely altered to assemblages of quartz, quartz–
associated with epithermal mineralization can be quite vari- alunite, and quartz–kaolinite where the rocks were permeated
able and often provides little exploration aid for vectoring by acidic fluids along fractures and breccia bodies. Some of the
toward mineralization (Simmons et al., 2005). However, cavities opened by acid attack are filled with native sulfur and
Leavitt and Arehart (2005), building on the earlier work by ore minerals, characteristically including pyrite, enargite, luzo-
Goldstrand and Schmidt (2000), conducted a detailed investi- nite, covellite, and gold-rich electrum. Barite and native gold
gation of hydrothermal alteration at Midas and showed that it are also common associates in cavities and in late veins.
could be used as an exploration aid. It is important to remem- The occurrence of many HS epithermal deposits at shallow
ber that the fluids causing much of the hydrothermal alteration depths above commonly Au-rich porphyry copper deposits
are not necessarily the same solutions that deposit the precious and the contemporaneity of the two (e.g., Arribas et al., 1995;
metals. For example, the precious-metal depositing fluid inter- Longo et al., 2010) clearly link the genesis of the two ore types,
preted for Midas would not have been particularly reactive as others have outlined (e.g., Sillitoe, 2010). There is also a
toward the wall rocks. However, if the heated groundwater clear kinship of this acid-sulfate epithermal ore type with many
entering the hydrothermal system at Midas had absorbed acid late-stage vein and breccia systems superimposed on porphyry
volatiles released from a degassing magma chamber at depth or copper deposits, such as the Butte, Montana, covellite–enargite
from a deeper boiling hydrothermal solution, then alteration ore with advanced argillic alteration (Meyer et al., 1968)
would be expected to be more intense. Indeed, Leavitt and and similar veins at El Salvador, Chile (e.g., Gustafson and
Arehart (2005) showed that hydrothermal alteration at Midas Hunt, 1975).
was apparently centered about the Colorado Grande and Gold
Crown veins, which are the principal structures. Leavitt and 13.15.2.3.1 Hydrothermal alteration
Arehart (2005) noted an extensive zone of weak propylitic The HS epithermal type ores are defined by the accompanying
alteration, consisting of calcite, chlorite, pyrite, and smectite, intense advanced argillic alteration, which at its most intense
which becomes more intense nearer the veins and where consists almost entirely of quartz. The rock may be up to 95 wt%
epidote appears as an alteration mineral. Closer to the veins, SiO2; all other primary constituents have been leached, leav-
smectite gives way to illite, the Fe:Mg ratio in chlorite increases, ing a spongy texture where original phenocrysts are removed.
pervasive K-alteration leads to adularia formation, and silicifi- The quartz-rich alteration is bordered by advanced argillic
cation is noted. Argillic alteration of wall rocks is common in assemblages always including alunite and kaolinite and
the upper parts of the Midas vein system, where it overprints variably including natroalunite, dickite, pyrophyllite, zunyite
earlier alteration events. Alteration halos surrounding the veins (Al13Si5O20(OH)18Cl), and diaspore. Further, kyanite and
extend a few tens of meters from the main veins, and the topaz are common in metamorphosed HS deposits like Brewer,
recognition of alteration patterns has aided in the exploration South Carolina, in the Carolina Slate Belt of SE USA.
in the district (Leavitt and Arehart, 2005). However, hydrother- The zoned alteration at Summitville, described by Steven
mal alteration varies with the lithology of the wall rocks and and Ratte (1960) and Stoffregen (1987), is clearly developed,
over time at Midas. providing a good example as a standard of comparison. The
 Geochemistry of Hydrothermal Gold Deposits 391

Summitville host rock is a quartz latite characterized by large hundreds of meters. The vuggy silica masses are all bordered
oligoclase and sanidine phenocrysts that, when leached out, laterally and vertically by alunite  kaolinite, dickite, and py-
yield a highly porous rock called ‘vuggy silica’ (Steven and rophyllite, all of which is capped by apparently steam-heated
Ratte, 1960) with centimeter-scale cavities, some of which are quartz–kaolinite alteration, which also contains chalcedony,
occupied by ore minerals. The near-vertical bodies of vuggy opal, sulfur, and rare alunite.
silica (Stoffregen, 1987) consist of quartz with small amounts The changes in rock composition upon intense hydrother-
of rutile, anatase, and residual zircon, all of which resist acid mal alteration have been well characterized in the Summitville
attack as shown by persistence of original concentrations of Ti, deposit. The bulk-rock compositions of altered Summitville
Zr, and Hf in the altered rock (see succeeding text). Vuggy silica quartz latite are depicted in Figure 8, which shows composi-
is surrounded on a scale of meters successively by alteration tions tabulated by Steven and Ratte (1960) and by Getahun
zones identified as quartz–alunite, quartz–kaolinite, illite– (1994), calculated on a volume basis (g cm3), to determine
kaolinite, and chlorite–montmorillonite (Figure 7). chemical gains and losses (Sales and Meyer, 1948). The data
In the quartz–alunite zone, phenocryst cavities are filled by show a complete leaching of Ca and Na in the quartz (vuggy
pink to white alunite in a groundmass entirely replaced by silica), quartz–alunite, and quartz–kaolinite zones and far
quartz. At the outer edges of the alunite zone, kaolinite joins into the montmorillonite–chlorite zone, reflecting destruction
alunite in the feldspar replacements and then becomes the of plagioclase. In the alunite zone, concentration spikes of
dominant Al-bearing mineral over widths of centimeters to a Na, Ba, K, Sr, Pb, and P reflect incorporation of those elements
meter. The kaolinite zone yields outward to illite-dominated into alunite and its analogues, such as woodhouseite (phos-
alteration, where feldspars are replaced by illite and quartz in a phate–sulfate). Titanium concentration is not changed in
tan to gray argillic groundmass. Further outward, in the illite– the altered rock, reflecting its residence in insoluble rutile and
montmorillonite zone, primary sanidine persists, but most pri- anatase. Zircon and Hf are similarly unaffected, owing to their
mary plagioclase is altered to illite–montmorillonite, as is the occurrence in refractory zircon. Iron is removed from vuggy
gray to tan groundmass. At a distance of many meters from the silica except where later pyrite fills porosity. Magnesium is
vuggy silica, the groundmass is green from montmorillonite and removed from vuggy silica and alunite zones where primary
locally chlorite, and although sanidine is unaltered, plagioclase biotite and hornblende are destroyed, but is apparently only
is clay-altered to various degrees. slightly affected in the kaolinite zone where it likely forms
In the Kasuga, Iwato, and Akeshi deposits of the Nansatsu colorless Mg-chlorite (Hedenquist et al., 1994). Magnesium
district, the minerals and zoning of alteration of the host and Ca are likely added to the rock in outer alteration zones
andesite are the same as at Summitville for all zones, but where carbonates form.
with added dickite, pyrophyllite, diaspore, and Mg-chlorite in Alumina is completely removed from quartz zone but is
the kaolinite zone at Iwato. Structurally, the alteration in the essentially unchanged in other zones, showing that alumina
Nansatsu ores lies in flattened subhorizontal amoeboid bodies from the vuggy silica zone is not remobilized to the kaolinite
on a scale of 100–200 m in diameter (Hedenquist et al., 1994; zone or laterally beyond. Gray and Coolbaugh (1994) found
Izawa and Cunningham, 1989) in contrast to the vertically slightly elevated Al2O3 on a weight basis in some kaolinite
elongated deposit at Summitville. zones, but not in others, although the weight basis precludes
The same minerals and zoning occur in other HS epithermal a clear conclusion regarding aluminum gains and losses.
deposits, including the giant gold deposits of the Yanacocha
district, Peru, where the tabular quartz-rich alteration zone
is gently dipping, 100–200 m wide, and broadly conforming 13.15.2.3.2 Trace elements and mineral associations
to volcanic layering, but locally cutting stratigraphy with a Primary ore minerals fill open space in intensely altered vuggy
subvertical attitude (Longo et al., 2010). Bordering the quartz silica where primary rock minerals were fully leached away
alteration, the sequence of 10–100 m wide alteration zones is before precipitation of the ore minerals. The common ore min-
again dominated by alunite, kaolinite, illite, smectite, and erals are enargite/luzonite, covellite, pyrite, gold-rich electrum,
chlorite–smectite. Similarly, at Pascua, Chile, Chouinard et al. and native gold. Description of electrum and gold textures
(2005a, 2005b) described masses of vuggy silica aligned along in relation to enargite/luzonite, covellite, and pyrite is rare, ap-
vertical structures and breccia but laterally coalesced over many parently owing to the fine-grained nature of gold and its incor-
poration into pyrite (see succeeding text). Assay data show a
positive correlation of Cu with Au in sulfide ore, although tex-
tures show that gold precipitated after enargite/luzonite in some
ores, such as Lepanto (Hedenquist et al., 1998), where the elec-
ill–mont chl–mont Fresh trum occurs with a postenargite assemblage including tennan-
ill–
qz–kal
a
qz– z

tite, chalcopyrite, sphalerite, and galena. Tennantite and

q

ka 2m
chalcopyrite are also common ore minerals; for example, at Sum-
mitville, Stoffregen (1987) found that the shallow enargite–
Figure 7 Cross section of the alteration envelope on a Summitville ore
covellite ores zone downward to tennantite–chalcopyrite, which
zone. Cu–Au mineralization is centered in the vuggy silica (qz), which is
bordered by zones of quartz–alunite (qz-al), quartz–kaolinite (qz-ka), were also in vuggy silica. In most settings, minor amounts of
illite–kaolinite (ill-ka), illite–montmorillonite, chlorite–montmorillonite, galena and sphalerite occur along with the copper minerals.
and fresh rock. Small rectangles represent occurrence of primary Vein or pore-filling barite with gold is a common late ore
sanidine. This schematic view is based on mapping of a pit exposure at form, cutting or overlapping gold ores dominated by pyrite or
the 3514 m elevation (M. Reed, unpublished data). enargite–pyrite. For example, Longo et al. (2010) describe late
392 Geochemistry of Hydrothermal Gold Deposits

qz–al ka–ill pyrite–tennantite–covellite–gold  barite  orpiment  sulfur

q qz–ka ill–mont mont–chl chl–mont fresh in unoxidized ore at Yanacocha, and at shallow depth, they
2.4
SiO2 report late gold intergrown with barite or quartz in vugs and
1.713
fractures. At Goldfield, Nevada, Vikre (1989) described ore
1.625 zones containing euhedral barite with quartz, pyrite, enargite/
1.3
luzonite, famatinite, bismuthinite, gold, and kaolinite, all
0.45
0.357 of which follow early quartz–pyrite–gold ore. Similarly, at
(0.356)
Paradise Peak, early quartz–pyrite–marcasite–sulfur breccia
Al2O3 with Au concentrations 1.8 ppm is cut by a black porous
0.00 quartz-matrix breccia containing barite, gold, silver, marcasite,
0.016
0.014 stibnite–bismuthinite, cinnabar, and sulfur but nearly lacking
0.011
Cu sulfosalts (John et al., 1989; Sillitoe and Lorson, 1994).
TiO2 Ore texture descriptions by Chouinard et al. (2005a, 2005b)

g cm−3
0.000 for Pascua, Chile, relate precipitation of gold, pyrite, and enar-
0.24
Fe gite together in an alunite–pyrite–enargite assemblage that fills
vuggy silica in a single ore stage. One distinctive pyrite textural
0.068 type with a dull luster and intergrown alunite hosts economi-
0.049
Vuggy silica

0.00 cally significant amounts of Au. The same pyrite contains in-
0.00
0.078 clusions of calaverite, gold, and electrum and is included within
enargite that also hosts grains of gold, electrum, gold tellurides,
0.032
CaO stibnite, cassiterite, and other phases containing Sn, Zn, Bi, Cu,
0.00 and Sb. These observations at Pascua establish contemporaneity
0.00
0.025 of Cu, Au, and As precipitation from a fluid in the enargite–
pyrite mineralization stage, as opposed to or in addition to a
0.010
MgO later stage characterized by tennantite–tetrahedrite and chalco-
0.00 pyrite, as at Lepanto (Hedenquist et al., 1998), or as a later stage
0.15
with barite, as outlined above.
0.100
(0.101)

K2O
0.00 13.15.2.3.3 Ore-fluid composition
0.1 0.094
Fluid inclusion homogenization temperatures and isotope frac-
(0.095)
tionation between minerals in HS epithermal Au deposits indi-
cate temperatures mostly in the range of 200–350  C at depths
Na2O
0.0 of about 200–600 m, inferred from coexisting vapor and liquid
0 5 10 15 20
inclusions in quartz and from geologic reconstructions of the
(a) Distance (m) paleosurface. At Julcani, Peru, sulfur isotope ratios for coexisting
pyrite (d34S ¼  1 to þ2%) and alunite (d34S ¼ þ 22 to þ25%)
40 indicate temperatures from 210 to 290  C with a mean of
260  C (Rye et al., 1992). At Pascua, coexisting S isotope ratios
21.05
in pyrite (d34S ¼ 3.4 to 5.3%) and alunite (d34S ¼ þ 15 to
Ba
0
þ20%) indicate temperatures ranging from 245 to 305  C
60
Sr

Figure 8 (a) Bulk chemical analyses (g cm3) of Summitville alteration

12.78
zones, plotted against distance (m) from the center of a vuggy silica zone
0
(left end). Circles show analyses from Getahun (1994). Squares show
Vuggy silica

7
g cm−3

5.02 analyses from Steven and Ratte (1960), but they are matched to
alteration types shown in circles, not to distance from vuggy silica.
Zr Dashed lines indicate composition of fresh quartz latite in the two
0 studies; where the two nearly overlap the numerical value of the one not
0.035
0.031 plotted is shown in parentheses beneath the other. Notice that fresh rock
(far right) MgO and CaO differ substantially in the two analyses, probably
reflecting unrecognized added carbonate in the Getahun (1994) samples.
Eu For other elements, the spread in the fresh rock compositions shown by
0.000
0.16
0.138
the vertical separation of the dashed lines indicates variation in parent
rock composition that provides a basis for understanding a likely
uncertainty in gains and losses in alteration owing to natural variation and
Hf to limitations of sample size. The strongly elevated concentrations of CaO
0.00
0 5 10 15 20 and MgO shown by squares at 17 m probably reflect addition of
(b) Distance (m) those elements in forming carbonates, whose presence is indicated
by 1.22 wt% CO2 in that sample. (b) Selected trace elements in the
Figure 8 (Continued) altered rocks shown in Figure 9(a).
 Geochemistry of Hydrothermal Gold Deposits 393

(Deyell et al., 2005). At Lepanto, Hedenquist et al. (1998) dominated by sulfuric acid formed from disproportionation
determined a range of 210–250  C for pyrite–alunite pairs. of magmatic SO2, according to the reaction:
Similar isotope findings, although less well constrained, for
1 þ 3
Summitville (Bethke et al., 2005; Rye et al., 1992) and the SO2 ðaq or gÞ þ H2 OðaqÞ ¼ 1 H ðaqÞ þ SO4 2 ðaqÞ
2 4
Nansatsu deposits (Hedenquist et al., 1994) indicate similar 1
temperatures. þ H2 SðaqÞ ! ðR1Þ:
4
A large number of measurements of liquid-rich fluid in-
clusions in the Nansatsu vein and vug-filling quartz, most of The products of the disproportionation reactions are vari-
which is texturally post-ore, have homogenization tempera- ably diluted by meteoric waters (Deyell et al., 2005; Rye et al.,
tures (Th) ranging from 160 to 240  C and most are accompa- 1992). The following Summitville isotopic data from Rye et al.
nied by vapor-rich inclusions (Hedenquist et al., 1994). Even if (1992) typify their findings for many deposits:
the Th values in the post-ore quartz underestimate mineraliza- 1. d34S (alunite) ¼ þ13 to þ25%. The heavy values distin-
tion temperatures, the coexisting vapor-rich inclusions indicate guish the sulfur in alunite as directly exsolved from a
boiling, yielding pressures corresponding to depths beneath magma, as opposed to derivation from oxidized H2S boiled
the water table ranging from 150 to 400 m. from a hydrothermal fluid, as is typical for steam-heated
At Julcani, Peru, Deen et al. (1994) measured a large num- acidic fluids. The heavy alunite reflects heavy sulfate sulfur,
ber of fluid inclusions in quartz, enargite, apatite, wolframite, which reflects isotope fractionation upon partitioning
and siderite. The homogenization temperatures for inclusions of
þ2% magmatic sulfur between light H2S and heavy
in enargite and ore-stage quartz largely overlap in the range of sulfate at a sulfide/sulfate ratio of 4 (2).
240–300  C. Homogenization temperatures for wolframite in- 2. The O and H in fluids, calculated from alunite: d18O ¼ 0 to
clusions cluster near 315  C. Deen et al. (1994) report no þ8%, and dD ¼  35 to  65%. The lighter H correlates
vapor-rich inclusions, indicating that maximum ore zone pres- with the lighter O, together indicating admixed meteoric
sures must exceed 100 bars at the high temperatures, indicating water with magmatic SO2 and H2O.
an open system paleodepth exceeding 1000 m. They also cite a 3. The O and H in fluids, calculated from wall rock kaolinite:
geomorphic reconstruction of the paleosurface indicating an d18O ¼ þ 4 to þ7%, and dD ¼  60 to 95%. The large
ore zone paleodepth exceeding 600 m. negative dD indicates meteoric-water dilution of fluids sim-
Fluid salinities in Julcani inclusions (Deen et al., 1994) ilar to those that produced alunite.
range from 17 wt% NaCl (equiv.) in early wolframite to
about 7 wt% in late stage and distal quartz and siderite. Ore- For the Nansatsu deposits, Hedenquist et al. (1994) provide
stage enargite lies between these in the range of 10–14 wt% analyses of pyrite, enargite, covellite, and native sulfur, all of
NaCl (equiv.). Deen et al. (1994) interpret these findings, in which occur in gold ore, wherein some vug fillings contain
combination with the isotopic evidence for meteoric-water enargite and pyrite together with gold. In the native sulfur
dilution of magmatic fluids, to indicate that the course of and sulfide minerals collectively, d34S ¼  5 to 0%, which
mineralization was accompanied by mixing of magmatic fluids Hedenquist et al. (1994) infer to indicate magmatic sulfur.
with meteoric waters. Similar conclusions for temperature and Alunite in these deposits has d34S ¼ þ 24 to þ34%, resembling
meteoric-water dilution are described by Hedenquist et al. the findings for Summitville, and together with the sulfide data,
(1998) for Lepanto on the basis of fluid inclusion and isotope indicates an aqueous sulfide/sulfate ratio of
3. A substantial
measurements, although salinities at Lepanto are in the range number of analyses of O and H isotopes in residual silica,
of 1–3 wt% NaCl (equiv.). alunite, and clays (mostly kaolinite, pyrophyllite, and Al-
Aside from Nansatsu, where abundant coexisting popula- chlorite) at Kasuga and Iwato yield ratios indicative of mixed
tions of liquid- and vapor-dominated fluid inclusion assem- magmatic and meteoric waters as inferred at Summitville. The
blages in late post-ore quartz indicate boiling (Hedenquist data also indicate that the Nansatsu clays formed with more
et al., 1994), most fluid inclusion studies reveal little boiling meteoric-water involvement than in the alunite alteration.
or occasional boiling in HS systems. For example, in a study
of late enargite–pyrite veins at Roşia Poieni, Romania, 13.15.2.3.5 Reaction of volcanic gas condensate with
Kouzmanov et al. (2010) identified coexisting liquid and quartz latite
vapor inclusions only in some quartz and pyrite, but in other Isotopic data combined with the open cavity textures observed
texturally distinguished pyrite and quartz, they found liquid in vuggy silica led Hedenquist et al. (1994) and Stoffregen
brine-only fluid inclusion assemblage with 35 wt% NaCl (1987) to suggest that magmatic gas condensed into ground-
(equiv.) and homogenization temperatures of 290–340  C. water at shallow levels, reacted with the host rock, and was
The boiling assemblages yield homogenization temperatures later followed by metal-bearing fluids that permeated the po-
of 250–284  C in fluids of
1 wt% NaCl (equiv.), correspond- rous rock, which was created by the acid leaching. To examine
ing to pressure on the order of 50 bars, which are temperatures the geochemical constraints on how the alteration pattern
and pressures similar to those in the Nansatsu deposits. exemplified by the Summitville system forms as a proxy for
the geochemical processes in most HS systems, a volcanic gas
13.15.2.3.4 Light-stable isotopes condensate was numerically reacted with the quartz latite of
The abundance of alunite and kaolinite facilitates the use of South Mountain in a computational model of alteration, fol-
isotope ratios of H, O, and S, useful in establishing the sources lowing the methods of Reed (1997) and using thermodynamic
and compositions of fluids responsible for hydrothermal alter- data described by Reed and Palandri (2006). The volcanic gas
ation in HS epithermal gold systems. HS ore fluids are is modified from the Augustine Volcano Spine 1D sample
394 Geochemistry of Hydrothermal Gold Deposits

analysis (Symonds et al., 1994) by a 5  dilution and an type and also reflecting a Ca enrichment they find relative to
increased Na2SO4 concentration, yielding a fluid with an ini- the primary rock. In other words, apparently some Ca leached
tial pH of 0.86 at 270  C and 55 bars. from plagioclase in the inner zones precipitates in the chlorite–
The initial fluid precipitates native sulfur, which is followed montmorillonite zone.
by quartz precipitation as the rock reacts with the evolving Across the alteration halo, the pH steps upward from 0.86
fluid (Figure 9(a)). With increasing amount of rock reaction, to 6.06, as the reaction with rock neutralizes the acid. In the
and thus decreasing water/rock ratio, the succession of key quartz and quartz–alunite intervals, aqueous Al3þ reaches a con-
alteration minerals following quartz is alunite, kaolinite, py- centration of 102.2 molal (160 ppm) at pH 1.0 (Figure 9(c)).
rite, barite, anhydrite, muscovite, hematite, chlorite, feldspars, Subsequent precipitation of kaolinite with rising pH removes
and calcite. This succession matches the observed sequence 99% of the aqueous Al3þ where pH has risen to 1.7. The high
(Steven and Ratte, 1960; Stoffregen, 1987) and can be broken solubility of Al-bearing and all other minerals except quartz at pH
into an outwardly zoned alteration series of quartz, quartz– less than 1.5 allows total leaching of rock forming elements from
alunite, quartz–alunite–kaolinite, quartz–kaolinite–muscovite the quartz alteration zone, yielding the highly porous vuggy silica
(illite), and propylitic (quartz, chlorite, feldspars, and calcite). that is subsequently filled by covellite, enargite, pyrite, and
The precipitation of calcite with the chlorite and feldspars electrum.
(Figure 9(b)) reflects its stability at neutral pH (5.7 at the ambi-
ent temperature), where its precipitation consumes CO2 from 13.15.2.3.6 Ore deposition
the magmatic gas. Steven and Ratte (1960) report whole-rock The ore minerals in most HS epithermal Au deposits filled
CO2 concentration of 1.22 wt% in the chlorite–montmorillonite open space or breccia matrices, fractures, or cavities in acid-
zone, reflecting carbonate minerals observed in that alteration leached wall rock, as opposed to replacement of wall rock

1
ka-ill
0 qz qz-al qz-al-ka Propylitic
Muscovite
Log mole/kg

Sulfur
-1 Anhydrite
tz Pyrite
Quar ite
-2 Alun

-3 Barite
Kaolinite
(a)

Albite
Calcite
-1 Chlorite
Log mole/kg

an20

-2
Microcline

an20

-3
Hematite
(b)

−pH
-1
SO42-
Log total molality

-2
Mg2+ Ca2+
-3
−pH
-4
K+
-5
-6 Al3+
-7
1000 100 10 1
(c) Water/rock
Figure 9 Computed reaction of a diluted volcanic gas condensate with quartz latite at 270  C. The water/rock ratio is determined from the initial mass
of water divided by the mass of rock. (a) Major minerals (abbreviation ‘an20’ refers to plagioclase with 20% anorthite component). Mineral alteration
zones corresponding to the Summitville alteration envelopes are marked, wherein muscovite is equivalent to illite, and the term ‘propylitic’ notably
includes the chlorite, feldspar and calcite assemblage shown in graph (b). (c) pH and total molalities of component species.
 Geochemistry of Hydrothermal Gold Deposits 395

along with alteration minerals. In HS systems, ore minerals which is written to the dominant species in the solution. The
may precipitate where hypogene fluids boil, conductively cool, equilibrium constant for this reaction increases by 13 orders
mix with cold groundwater, or mix with a reactive fluid, such as of magnitude between 300 and 100  C, largely owing to the
an acidic and Fe-rich magmatic condensate residual from the increased intrinsic stability of gold with decreasing tem-
wall rock alteration stage (as above). An exception to open perature; thus, simple temperature decrease plays a key role
space-filling ore is the stage I ore at Pueblo Viejo, Dominican in precipitating electrum. Despite the much larger aqueous Ag
Republic (Kesler et al., 1981; Muntean et al., 1990), where Au concentration than Au (Figure 10(c)), the computed electrum
and pyrite precipitated together upon sulfidation of wall rock is gold-rich (Figure 10(b)), as observed in the natural setting,
iron by the hydrothermal fluids, as occurs in orogenic gold because the gold end-member is intrinsically quite stable rela-
deposits (e.g., Reed, 1997). tive to silver, and the temperature decrease has a stronger effect
The measurements of fluid inclusion salinities and isotopic on the gold reaction than it does on the corresponding silver
ratios of H and O demonstrate that boiling and meteoric-water reaction involving aqueous Ag chloride.
dilution of ascending magmatic fluids are active in HS epithermal The effects of cooling and acidic conditions on ore mineral
ore-forming systems. These processes were evaluated in a series precipitation, as outlined earlier in the text for the cold water
of numerical chemical equilibrium model calculations of simple mixing, apply similarly to boiling and to mixing with an acidic
cooling, boiling, cooler water mixing, and acidic water mixing, Fe-rich magmatic condensate residual from the wall rock alter-
each one starting from a magmatic fluid at 325  C. The starting ation (Figures 11 and 12), both of which were computed for
fluid is derived, in part, from fluid inclusion compositions and the same starting fluid used for cold water mixing. Boiling
equilibrium constraints applied in a model of wall rock alteration yields covellite and electrum, but little pyrite until tempera-
in the magmatic-hydrothermal system of Butte, Montana (Reed tures cool below 175  C (Figure 12). Electrum precipitates at a
et al., 2013). The starting fluid Au aqueous concentration is at the higher temperature upon boiling than in cold water mixing
low end of the range (0.1–5.0 ppm Au) found by Kouzmanov because a large quantity of aqueous H2S is displaced to the gas
et al. (2010) for fluid inclusions in enargite and pyrite at Roşia phase, driving reaction R3 to the right, precipitating gold. The
Poieni, Romania. The results of three of the model runs displacement of S to the gas phase also precludes precipitation
(Figures 10–12) illustrate the essential chemical processes. of native sulfur in the mineral assemblage.
Progressive mixing of the 325  C starting fluid with 5 kg of Mixing of the magmatic fluid with a 100  C residual acidic
25  C meteoric water causes quartz, enargite, covellite, pyrite, Fe-rich fluid from the wall rock reaction calculation in the
sulfur, Au-rich electrum, stibnite, proustite, galena, and sphal- preceding text (at a water/rock ratio of 36) yields abundant
erite to precipitate (Figure 10(a) and 10(b)), as pH changes pyrite with covellite and smaller amounts of enargite and
from 2.5 at 325  C to 2.1 at 75  C (Figure 10(c)). Sulfide alunite (Figure 12) at high temperature. Pyrite and covellite
precipitation is driven by several factors (Reed and Palandri, are joined by sulfur, stibnite, electrum, and barite at lower
2006) that function despite the acidic pH: (1) decreasing tem- temperature, but not by galena and sphalerite, which form
perature drives metal chloride complex dissociation; (2) dilu- upon cold water mixing (Figure 10(a)). The pH changes
tion reduces the concentration of metal-complexing chloride; from 2.5 at 325  C to 0.9 at 135  C. The more acidic pH at
and (3) the sulfide minerals become intrinsically less soluble at low temperature prevents precipitation of galena and
lower temperature. The acidic pH favors pyrite plus covellite sphalerite; however, the added barium from the wall rock
and enargite over the lower-sulfur alternatives, chalcocite (and fluid enables barite to form with the low-temperature ore
bornite and chalcopyrite) and tennantite, illustrated by the assemblage. Similarly, aluminum in the wall rock-derived
following reaction (see Reed and Palandri (2006) fluid forms alunite mixed with pyrite and covellite at high
temperature. The low pH in this system results in substantial
1 1
2 CuSðcvÞ þ H2 OðaqÞ ¼ Cu2 SðccÞ þ SO4 2 ðaqÞ þ Hþ ðaqÞ sulfur precipitation, more than for cold water mixing or boil-
4 2 ing, which produced none.
3
þ H2 SðaqÞ ! ðR2Þ The ore deposition calculations illustrate that cooling, cold
4
water mixing, boiling, and acidic water mixing all yield the
for more details) that shows increased Hþ activity favors covel- observed ore mineral assemblages. Given the wide variation in
lite (cv) over chalcocite (cc). the details of natural assemblages and their relative timing as
In the cold meteoric-water dilution, aqueous Auþ is com- outlined above and the similarly various computed assem-
plexed almost entirely by AuHS, which holds a minimum of blages, depending on the processes and the details of fluid
85% of the Au; AuCl2 and AuCl32 contain the remainder compositions, it is likely that mixing, boiling, and cooling all
(Figure 10(d)). Strong dominance of the Au bisulfide applies play roles in HS epithermal Au systems over time. In all of the
over all of the acidic pHs and temperatures up to at least ore precipitation processes, Pb, Zn, and most of the Ag remain in
350  C, despite the much greater concentration of H2S than the solution at temperatures above 200  C (e.g., Figure 10(d)),
HS (Figure 10(d)) owing to the acidic pH. The large aqueous thus providing for their transport to a distal setting, where they
sulfide concentration in the HS ore fluids is sufficient, despite may precipitate in base metal–silver veins.
the low pHs, to assure enough HS to form AuHS instead of In summary, HS epithermal Au deposits form at tempera-
chloride complexes. Thus, gold precipitation is controlled by tures of 240–340  C, where magmatic SO2 disproportionates
the following gold precipitation reaction: to yield sulfuric acid and H2S that mix with meteoric water in a
subvolcanic environment at depths of 200–1000 m. The sulfu-
8 AuHS ðaqÞ þ 4 H2 O ðaqÞ ¼ 8 Au þ 7 H2 S ðaqÞ ric acid fluids leach all of the major constituents but silica from
þ HSO4 ðaqÞ þ Hþ ðaqÞ ! ðR3Þ the host rock, thereby creating porosity that is subsequently
396 Geochemistry of Hydrothermal Gold Deposits

(a)
−2 S⬚
Quartz
sl cv en
−3 py
gn
−4 en
Log mole kg-1 kg-1 S⬚ orp

−5

−3
−4 (b) Pyrargyrite Stib
Proustite
−5 Au
−6 Ag
−7
−8
−9
(c)
−1 Na+ HS−
SO42− −pH
−2
Log total molality

−3
Zn2+ Pb2+
−4 Ag+
−5 Fe2+ Sb(OH)3
−6 Au+
−7
−8 Cu+ H2AsO3−

(d)
−2
H2S CuCl2−
−3
CuCl
−4
Log molality

−5 AgCl2−
AuHS
−6
HS−
−7
−8 AuCl2− AgHS
−9 AuCl32−

75 100 125 150 175 200 225 250 275 300 325
Temperature (⬚C)
Figure 10 Model results for mixing of 325  C magmatic fluid with 25  C water with consequent temperature decrease, dilution, and reaction. At the low
temperature end, 5 kg of cold water had been added to 1 kg of magmatic fluid. (a and b) Log of rate of mineral precipitation in units of moles per
kilogram of initial water per kilogram of admixed cold water. (c) Total molalities of aqueous components. For Naþ, the downward slope of the curve with
decreasing temperature shows the effect of dilution alone on concentration. The steeper slopes of other curves (e.g., Cuþ, Fe2þ, and Auþ) reflect
precipitation of those elements in minerals. (d) Molalities of sulfide species and the dominant individual complexes of Cu, Au, and Ag with HS and Cl.
Chlorides of Cu and Ag and a bisulfide of Au dominate. Mineral abbreviations are as follows: al, alunite; bar, barite; cv, covellite; en, enargite; gn, galena;
orp, orpiment; py, pyrite; pyro, pyrophyllite; sl, sphalerite; and stib, stibnite.

permeated by copper- and gold-bearing fluids that precipitate chalcopyrite, indicates acidic conditions in the ore precipitation
enargite/luzonite, covellite, pyrite, and electrum in the pore zone, but not so acidic as to prevent sphalerite, galena, and
space. The metal-bearing fluids are apparently shallow deriva- tennantite–tetrahedrite from forming in some settings.
tives of the same fluids that produce porphyry copper miner- Although the acidic conditions are probably partly residual
alization at depth, as argued by Hedenquist et al. (1998) based from the earlier period of acid-sulfate leaching, it is likely that
mainly on their findings on the HS epithermal ores of the the metal-bearing fluid is also acidic owing to its derivation by
Lepanto system and the subjacent Far Southeast porphyry copper cooling of a metal-bearing porphyry copper fluid as explained
deposit, Luzon, Philippines. Copper and gold precipitate where by Reed et al. (2013), who propose that a late aliquot of the
fluids cool upon boiling, mixing with meteoric waters, or mixing same fluid that forms the potassic and sericitic alteration in
with residual acid-sulfate fluids in the porous host rock. The porphyry copper deposits likely also forms the vein deposits of
prevalence of covellite, pyrite, and enargite, as opposed to pyrite, covellite, enargite, and chalcocite in Butte and similar
 Geochemistry of Hydrothermal Gold Deposits 397

Boiling
−3
−4 py Quartz
−5 cv

Log mol kg−1 ⬚C

Proustite
−6
−7 Stib Au
−8
Stib Pyrargyrite
−9
Ag
−10
−11

100 125 150 175 200 225 250 275 300 325
Temperature (⬚C)
Figure 11 Model results of mineral precipitates from boiling of a 325  C magmatic fluid from 325 to 100  C, expressed as the rate of mineral
precipitation per degree per kilogram of initial water. See caption to Figure 10 for mineral abbreviations. The ore mineral assemblage is dominated by
covellite and electrum at high temperature, but little enargite and pyrite and no sulfur. Electrum precipitates at a higher temperature (275  C) upon
boiling than it does by mixing, reflecting the removal of sulfur to the gas phase, which liberates Au from AuHSo. The pH remains acidic throughout the
run, from 2.4 at 325  C to 1.2 at 100  C, which, in combination with sulfur loss to the gas phase, precludes precipitation of galena and sphalerite at low
temperature, in contrast to cold water mixing effect (Figure 10).

Mixing acidic fluid

−1 S⬚ py
−2 pyro
Proustite Quartz
−3
Log mol kg−1 kg−1

Pyrargyrite bar cv
−4
Stib en
−5 al
−6 Au
−7
−8 Ag
−9
Stib

−10
150 175 200 225 250 275 300 325
Temperature (⬚C)
Figure 12 Model results of mineral precipitates from mixing of a 325  C magmatic fluid with an acidic and Fe-rich 100  C diluted magmatic gas
condensate that had reacted with quartz latite wall rock (Figure 9). The ordinate units are log moles of mineral per kilogram of initial water per kilogram
of admixed acidic water. See caption to Figure 10 for mineral abbreviations. The pyrophyllite and alunite at high temperature reflect Al leached from
altered rock. This run forms more sulfur than the others, reflecting the more acidic conditions, and electrum precipitates at a lower temperature,
reflecting the effect of acid in displacing reaction R3 to the left.

deposits. The particular qualities of pH and metal concentra- epithermal transition’ should be evaluated using a system-
tions of that fluid depend on its ascent history, including extent wide approach. The recognition by Heinrich et al. (1999)
of wall rock reaction, and its P–T path with or without conse- that a low-density, gold-rich magmatic fluid separates from a
quent phase separation. When that fluid arrives in the shallow much denser magmatic fluid has led to a shift in understanding
environment, it may boil, mix with groundwater, or cool by the geochemical processes related to how magmas release
conduction, yielding variations in the ore mineral assemblages metals to the much shallower epithermal systems (Heinrich
and variations in the relative timing of assemblages. The acidic 2005; Heinrich et al., 2004; Williams-Jones and Heinrich,
conditions that define the ore type, however, always yield 2005; see also Chapter 13.1). Thus, the hypothesis of vapor-
advanced argillic alteration surrounding ores that include py- phase transport of precious metals by supercritical, low-density
rite, covellite, enargite, sulfur, and gold. magmatic fluids offers an attractive mechanism for the forma-
tion of LS epithermal deposits, but further investigations of
this hypothesis, including laboratory experiments, are needed.
13.15.2.4 New Ideas about the Geochemistry of
New Pb and Re–Os isotopic studies on ore minerals, such
Epithermal Deposits
as electrum (Kamenov et al., 2007; Saunders et al., 2011c;
In the past few decades, it has been increasingly apparent that Figure 13), provide evidence for a deep magmatic source of
there is a genetic link between some epithermal ores and gold in LS ores in the Great Basin of the USA and that mafic
porphyry ore deposits (e.g., Hedenquist and Lowenstern, magmas are the most likely source of the precious metals
1994; Saunders, 1991; Saunders and May, 1986; Sillitoe, (Hames et al., 2009; Kamenov et al., 2007; Saunders et al.
1997), and thus the geochemistry of the ‘porphyry to 2008). If vapor-phase transport of metals and metalloids are
398 Geochemistry of Hydrothermal Gold Deposits

0.40 (2011) and Berger and Henley (2011) discuss the process of
magmatic-vapor expansion and the formation of HS gold min-
eralization at El Indio, Chile. More research on the role of
vapor transport of metals and nanoparticles in hydrothermal
0.35
16 Ma ore formation appears to be warranted, including laboratory
SLEEPER experiments to test this hypothesis. However, for now, one
Electrum possible implication is that perhaps epithermal ore-forming
0.30 15 Ma processes are not restricted to the shallow crust as has been
NATIONAL
initial

previously assumed.
Electrum
187Os/188Os

0.25
13.15.3 Carlin-Type Gold Deposits
13.15.3.1 Introduction
0.20 14 Ma, Wanapum CTD are epigenetic, disseminated, auriferous pyrite deposits
16 Ma, Grande
Ronde-R1 enriched in As, Sb, Hg, and Tl that are typically hosted in
calcareous sedimentary rocks. They are named after a deposit
0.15 near Carlin, Nevada, that spurred exploration for similar de-
posits around the world. In these deposits, gold ore commonly
17 Ma, Imnaha occurs at intersections between faults and permeable reactive
strata typically below an impermeable caprock and generally
0.10
exhibits a central high-grade zone of carbonate dissolution and
18.7 18.8 18.9 19.0 19.1 argillic alteration with micron-sized disseminated pyrite. Per-
206Pb/204Pb vasively silicified rock, called jasperoid, is generally present
within or near ore. In unsilicified rocks, many open fractures
Figure 13 Plot of the initial Os and Pb isotopic composition of Sleeper and pores are often lined with macroscopic crystals of orpi-
and National (NV) electrum (data are from Kamenov et al. (2007) and ment, realgar, and calcite. In jasperoid, most fractures and vugs
Saunders et al. (2011c)) relative to the range of values from Columbia
are lined with drusy quartz and may contain kaolinite, stibnite,
River basalt flows Wanapum, Grande Ronde, and Imnaha (data from
barite, or calcite (Cline et al., 2005).
Chamberlain and Lambert, 1994; Chesley and Ruiz, 1998).
The three largest concentrations of CTD are in the western
USA and in two areas in southern China (Figure 14). The
amount of gold present in CTD in north-central Nevada
important for the formation of LS epithermal ores, then the (5500 t) is enormous and is exceeded only by the Witwatersrand
volatility of these elements should be considered and that in South Africa. The CTD in the Dian-Qian-Gui area and West
too is only recently being evaluated (Saunders et al., 2008; Qinling belt in southern China each contains about an order
Saunders and Brueseke, 2012). Lindgren (1933) and Saunders magnitude less gold (400 t), but is not as thoroughly explored.
(1990, 1994) showed from the interpretation of electrum and Gold ore is mined by both open-pit and underground
silica textures that gold colloids played an important role in operations, with lower ore grades in the former (less than
forming some bonanza LS epithermal ores such as at the 1 g t1) and higher ore grades in the latter (more than
National deposit in Nevada. Saunders (1994) and Saunders 34 g t1). Because most of the gold resides in pyrite, it must
and Schoenly (1995) proposed that the electrum colloids form be oxidized in a roaster or autoclave to recover gold with
due to boiling of the ore-forming fluids. The new isotopic data cyanide solutions. Natural weathering and oxidation of the
of Kamenov et al. (2007) and Saunders et al. (2011c; Figure 13) pyritic ores eliminates this step and permits direct recovery of
led to a provocative new hypothesis; it can be argued that gold, gold by cyanide heap leach methods.
silver, and copper nanoparticles, which are smaller than col- The geochemistry of CTD ore formation, described herein,
loids, nucleated at much deeper levels than the epithermal is fairly well understood. In several districts, chemical and
environment, locked in a primitive isotopic signature in the isotopic data constrain the source of ore fluid components
process, and traveled to the site of epithermal ore formation as and the age of mineralization enabling development of genetic
nanoparticles (Kamenov et al., 2007; Saunders, 2012; Saunders models that relate mineralization to the geotectonic evolution
et al., 2008, 2011a, 2011b). In this new and relatively untested of the region. However, the evidence collected from different
hypothesis, the same low-density supercritical Cu–Au–Ag-rich districts permit multiple interpretations regarding the source of
fluid found by Heinrich et al. (1999) is interpreted to be close gold (magmatic or sedimentary) and the role of magmatic,
to saturation with respect to electrum, naumannite (Ag2Se), metamorphic, and meteoric fluids in gold transport and depo-
and chalcopyrite and that perhaps cooling from
900 to sition. Thus, no single model adequately explains the origin of

700  C leads to saturation or supersaturation with respect every district.
to the aforementioned minerals and they form nanoparticles at Recent models for Nevada CTD call upon magmatic
a relatively high temperature and deep setting. Recently, (Muntean et al., 2011) or sedimentary (Large et al., 2011a,
Kouzmanov et al. (2010) found indirect evidence of solid 2011b) sources for gold and transport to the site of ore depo-
precious-metal nanoparticles trapped in fluid inclusions in sition by magmatic or metamorphic/meteoric fluids, respec-
enargite from a HS epithermal system. Further, Henley et al. tively. Evidence from CTD in the West Qinling belt suggests
 Geochemistry of Hydrothermal Gold Deposits 399

Figure 14 Location of Carlin-type deposits (CTD) (black or red dots) in the Great Basin, western USA, and in the West Qinling belt and
Dian-Qian-Gui area of southern China.

they may be epizonal analogues of nearby mesozonal orogenic rocks that were either thrust over the platform margin during
gold deposits with gold transport by metamorphic fluids the early Mississippian Antler orogeny or overlain by Mississip-
(Hofstra et al., 2005; Mao et al., 2002). Likewise, evidence pian shales deposited in the adjacent foredeep (Hofstra and
from CTD in the Dian-Qian-Gui area is consistent with an Cline, 2000). In the West Qinling belt, orogenic gold deposits
epizonal orogenic model (Hofstra et al., 2005; Hu et al., 2002; occur in metamorphosed Devonian and Carboniferous siliciclas-
Su et al., 2009a, 2009b). If correct, then one might view the so- tic rocks of the accretionary prism, and CTD occur in Triassic
called CTD in China as simply epizonal orogenic gold deposits Songpan–Ganzi basin flysch or Neoproterozoic–Permian
that have developed in a specific tectonostratigraphic setting. miogeoclinal rocks along the northern margin of the Precam-
brian Yangtze craton that were deformed by a south-verging
foreland fold-and-thrust belt (Mao et al., 2002). Both gold
deposit types are inferred to have formed after peak deformation,
13.15.3.2 Age and Geologic Setting of CTD
metamorphism, and magmatism (c.245–220 Ma) during the
Many, but not all, CTD in Nevada are essentially the same age final stages of contraction (Mao et al., 2002), transpression
and cospatial with late Eocene dikes and thermal halos, aero- (Vielreicher et al., 2002), or postcollisional uplift and extension
magnetic anomalies, and concealed subduction-related mag- of the Triassic Qinling collisional orogen. In contrast, in the
matic centers (Ressel and Henry, 2006), which were emplaced Dian-Qian-Gui area, CTD are hosted in Cambrian to Middle
during slab rollback in an extensional tectonic regime. However, Triassic age platform carbonate (terrestrial basin), transitional,
the CTD generally lack, or fail to fit into, alteration zonation and basinal siliciclastic rocks, and locally in mafic intrusions
patterns typically developed around epizonal porphyry intru- (diabase and gabbro) or Late Permian Emeishan volcaniclastic
sions (Cline et al., 2005). They generally occur in Paleozoic rocks along the southwest margin of the Yangtze craton. Field
carbonate rocks deposited near the platform margin (Cook, relationships and unconventional isotopic dates suggest the
2005) along reactivated high-angle fault systems that may be deposits are Cretaceous in age (Hu et al., 2002; Su et al., 2009a,
inherited from Neoproterozoic rifting and formation of the 2009b), although a new Re–Os isochron date of 193 13 Ma on
western continental margin of North America (Lund, 2008). arsenian pyrite and a 40Ar–39Ar date of 194.6  2 Ma on sericite
These faults were reactivated during subsequent extensional from Jinfeng both indicate an older age (Chen et al., 2007;
and contractional orogenies and served to localize Paleozoic Gu et al., 2010).
hydrothermal dolomite and sedimentary exhalative deposits
(SedEx) of Zn, Ba, and Au mineralization, as well as Mesozoic
13.15.3.3 Ore and Gangue Minerals
and Cenozoic plutons and related hydrothermal ore deposits
(Emsbo et al., 2006). CTD typically occur below a less In all CTD, there are pre-ore hydrothermal breccias, followed
permeable caprock consisting of shaley slope or basin facies by a main ore stage where fluids dissolve or replace minerals in
400 Geochemistry of Hydrothermal Gold Deposits

the host rocks, and subsequently there is a late ore stage when illite, illite–smectite, smectite, and chlorite. Pre-ore carbonate
changes in temperature, fluid mixing, or phase separation minerals are generally dissolved or replaced during the main
(boiling) cause minerals to precipitate in open fractures and ore stage, but hydrothermal calcite, dolomite, ankerite, and
vugs (Figure 15). siderite also precipitated during main stage ore formation.
Most of the gold in the ores resides in pyrite that replaces Vug or fracture-filling minerals of the late ore stage include
Fe-bearing minerals in the host rocks and commonly forms pyrite, marcasite, orpiment, realgar, native arsenic, stibnite,
rims on pre-ore pyrite (Figure 16). Hence, pyrite is the princi- cinnabar, lorandite (TlAsS2), galkhaite [(Cs,Tl)(Hg,Cu,
pal ore mineral in CTD. Native gold is quite rare and generally Zn)6(As,Sb)4S12], sphalerite, tellurides, quartz, dickite, calcite,
is only visible in supergene goethite pseudomorphs of main barite, and adularia (Figure 17). In a few deposits, millimeter-
ore-stage pyrite. In some deposits, gold also occurs in marcasite sized native gold is present in vugs or calcite veinlets. Micron-
or arsenopyrite, jasperoidal quartz, and organic carbon. Main sized gold has been observed in late ore-stage marcasite and
ore-stage silicate minerals include quartz, dickite, kaolinite, cinnabar.

13.15.3.4 Geochemistry of Rocks and Pyrite

The application of multivariate statistical techniques, such as
factor analysis or cluster analysis, to whole-rock multielement
Decalcified ore geochemical data has proved useful in the identification of
element associations of geologic significance (Hofstra and
Cline 2000; Hofstra et al., 2011). Element associations for
rock samples from CTD determined by these methods often
correspond to the marine carbonate (CO2, Ca, Sr, Mg, and
Mn), terrigenous detritus (Al, Ti, K, REE), and organic (V, Ni,
Mo, Se, and Zn) components of the host sedimentary rocks.
Calcareous waste For example, Figure 18 profiles the correlation of organic
carbon, organic component factor scores, key elements,
and the @ 34Ssulfide composition of the Popovich Formation,
which is an important host for gold in the northern Carlin
Figure 15 Photograph of decalcification front in Ordovician–Silurian Trend. Carlin-type element associations identified in this
Hanson Creek Limestone, Jerritt Canyon. Pre-ore calcite veins in manner include silicification with stibnite (Si and Sb) or
limestone are absent in decalcified limestone. Fractures in ore are filled precipitation of main ore-stage pyrite and late ore-stage
with orpiment and realgar. Pen for scale.
orpiment, realgar, cinnabar, and lorandite (Au, As, Sb, Hg,
and Tl). Depending on the pre- and post-ore history of the
district, element associations due to other igneous and hydro-
thermal events may also be apparent, such as SedEx minerali-
zation, felsic or mafic dikes, contact metamorphic aureoles,
or base metal mineralization. Such differences are important
and make it possible to distinguish preexisting geochemical
suites from those imposed by Carlin-type hydrothermal sys-
tems. Likewise, multivariate statistical analyses of micro-
analytical scans (e.g., LA-ICP-MS) across pyrite present in the
ore can aid identification of the trace element associations
present in diagenetic pyrite (V, Ni, Co, Cu, Mo, and Se) and
main ore-stage pyrite (Au, As, Sb, Hg, Tl, (Cu, Pb, Te)) (Cline
et al., 2005; Large et al., 2009).
To identify elements that were immobile, introduced, or
depleted and to measure mass losses or gains associated with
alteration and mineralization, the whole-rock multielement
compositions of barren, unaltered host rocks and ore-grade
equivalents can be compared on an isocon diagram (Grant,
1986; Hofstra and Cline, 2000). An isocon diagram for the
10 microns
Carlin deposit (Figure 19(a)) shows that SiO2, Al2O3, Fe2O3,
K2O, TiO2, and P2O5 were immobile; CaO, CO2, MgO, Na2O,
and MnO were depleted; and S, As, Sb, Tl, Hg, and Au were
introduced. The concentration of immobile elements in the
ore is enriched by as much as 50% due to mass loss associated
Figure 16 Reflected light photomicrograph of zoned euhedral pre- with dissolution of carbonate minerals and argillization of
ore-stage pyrite with rim of ore-stage pyrite, Deep Star deposit, Carlin albite in the silty laminated limestone host rock. Where this
Trend. Ore-stage pyrite is more fine grained, anhedral, and darker than approach has been properly applied to other CTD, evidence
pre-ore pyrite. for analogous gains and losses is generally detected (e.g.,
 Geochemistry of Hydrothermal Gold Deposits 401

Calc
ite

Orpiment Barite
Realgar

Orpiment Cinnabar
Marcasite
(a) (b)

Kaolinite

Stibnite

Quartz
(c) (d) (e) Kaolinite

Figure 17 Images of late ore-stage minerals: (a) orpiment, realgar, and calcite in decalcified limestone (Getchell); (b) orpiment, marcasite, and
cinnabar cemented by barite, Jerritt Canyon; (c) Stibnite rosette, Jerritt Canyon; (d) Drusy quartz and kaolinite, Deep Star (Thompson, 2005); (e) SEM
image of kaolinite, Deep Star mine.

500
Org. C Mo V Zn

Factor 2 BSR-
600 UM open

SD
700

Ji
800 SD
Feet

PL BSR-
900 Ji
PL
SO4 limited
WSPL

1000 Ji

WSS
1100

1200
0 1 2 3 4 5 6 4 40 400 −15 −10 −5 0 5 10 15
Factor score or wt% ppm d 34S ‰

Figure 18 Downhole profile of metals and @ 34Ssulfide in strata (UM upper mud unit, SD soft sediment unit, PL planar unit, WSPL upper wispy planar
subunit, and WSS lower wispy unit) of the Devonian Popovich Formation, from Hofstra et al. (2011). Factor 2 element suite comprised Ni, Cu, V, Zn, Cd,
organic C, Mo, Y, and Cr. While Se was not determined, studies by Large et al. (2011a, 2011b) indicate it is part of the metalliferous black shale element
suite. Sulfur isotope data are suggestive of bacterial sulfate reduction (BSR) in open and sulfate-limited marine environments. Intervals intruded by
Jurassic dikes (Ji) are masked in gray.
402 Geochemistry of Hydrothermal Gold Deposits

1 000 000
SiO2

Al2O3
Fe2O3
Isocon
LOI

K2O
100 000 Constant Al, K, Fe, Ti, P
50% mass decrease

CO2
CaO
MgO

P2O5
TiO2
S
10 000

As C
Ba
1000 Hg
Cr
Au Na2O
100 TI
Altered
MnO
Sb
10
Mo
1

.1

Zone 3, N = 5
.01 Constant mass Zone 2, N = 4
Zone 1, N = 4
1E-3
1E-3 .01 .1 1 10 100 1000 10 000 100 000 1 000 000

Unaltered
(a) N = 14

1 000 000
Isocon Si
100 000 Constant Al, K, Fe, Ti Al
62% mass decrease Fe K
S Ca
10 000 Mg
As Ti
1000 P
Sb
Hg Ba
Altered limestone

Cr
Au Ni Zn V B
100
TI Cu Mn Sr
ppm

Se
10

Ag Co
1

0.1

0.01 Constant mass

0.001
0.01

0.1

10

100

1000

10 000

100 000

1 000 000
0.001

ppm
(b) Unaltered limestone

Figure 19 Isocon diagrams showing element mobility in (a) Silurian–Devonian Roberts Mountains Fm. at Carlin (data from Bakken, 1990) and
(b) Ordovician–Silurian Hanson Creek Fm. at Jerritt Canyon (data from Hofstra, 1994). Elements that plot above or below the isocon are, respectively,
introduced or depleted. In each deposit, the immobility of Fe and the introduction of S and Au are clear evidence of gold precipitation by sulfidation
of host-rock iron.

Jerritt Canyon, Figure 19(b)). The immobility of Fe and the et al., 1991), discussed further in the succeeding text. Although
introduction of S, Au, As, Sb, Hg, and Tl, evident on the isocon less common, isocon diagrams for samples collected from
diagrams, require that the Au- and trace element-rich pyrite in the margins of a few deposits show that Fe locally was intro-
the ore precipitated where H2S in ore fluids reacted with duced, which suggests ore fluids locally mixed with external
Fe-bearing minerals in the host rocks (sulfidation; Hofstra Fe-bearing fluids (Cail and Cline, 2001; Kesler et al., 2003).
 Geochemistry of Hydrothermal Gold Deposits 403

In many early studies of CTD, multielement analyses were generally are enriched more than one hundred times their
not conducted, and only a small suite of trace elements Clarke values (Au, As, Sb, Hg, and Tl), whereas those that
were determined on ore-grade material. In such cases, an alter- form chloride complexes (Ag, Pb, Zn, and Cu) or oxyanions
native approach is to scale the trace element concentrations in (Mo and W) are not as enriched (<100  Clarke). The strong
the ore (e.g., >1 ppm Au) to their average concentrations in the enrichment of the so-called Carlin suite of trace elements in the
crust (Clarke values; Levinson, 1974). Element enrichment deposits is a direct reflection of the importance of sulfide
plots constructed with such data from 31 CTD in Nevada complexing in ore fluids. Trace element data from CTD in the
(Figure 20) show elements that form sulfide complexes West Qinling belt and Dian-Qian-Gui area of southern China

100 000
Nevada
10 000 CTD

1000
Enrichment factor

100

10

0.1

(a) 0.01
Tl Hg Sb As Au Ag Pb Zn Cu Mo W

100 000
D-Q-G
10 000

1000
Enrichment factor

100

10

0.1

(b) 0.01
Tl Hg Sb As Au Ag Pb Zn Cu Mo W
100 000

10 000 W Qinling

1000
Enrichment factor

100

10

0.1

0.01
(c) Tl Hg Sb As Au Ag Pb Zn Cu Mo W
Figure 20 Trace element enrichment plots for ore (1 ppm Au) relative to average crust in (a) 31 Nevada Carlin-type deposits (CTD) from each
significant gold district. In (b) and (c), data from seven CTD in the Dian-Qian-Gui area and five CTD in the West Qinling belt (colored diamonds) are
shown relative to data from CTD in Nevada (black bars). The results suggest that the chemistry of ore fluids in each area was very similar.
404 Geochemistry of Hydrothermal Gold Deposits

(a) (d) At the large (220 t Au), high-grade (24.7 g t1) Meikle de-
posit in the Carlin Trend of Nevada, evidence of carbonate
cm
dissolution and sulfidation is exceedingly clear because the
ore occurs in an unusual ferroan dolomite host rock (Emsbo
et al., 2003; Figure 21(a) and 21(b)). LA-ICP-MS microanaly-
ses of ore pyrite that replaced ferroan dolomite (Figure 21(b)
and 21(c)) provide a clear indication of the relative concentra-
tions of trace elements (As > Sb > Au > Cu > Tl > Ag > Hg > Pb;
(b) (c)
Figure 21(e) and 21(f)). Microanalyses of ore pyrite from
cm cm other CTD show that its composition varies from deposit to
deposit, both within and between different mining districts.
For example, a greater abundance of Te and Cu is present in ore
pyrite from certain deposits and districts (e.g., Barker et al.,
2009). The sulfur isotopic composition of ore pyrite in CTD
also varies. For example, in the northern Carlin Trend, the @ 34S
of ore pyrite from the high-grade Meikle deposit and the cen-
tral high-grade part of the enormous Betze-Post deposit is
about 10% (Emsbo et al., 2003), whereas values in the
Screamer section on the margin of Betze-Post are near 0%
(Kesler et al., 2005; Hofstra et al., 2011; Figure 22). Likewise,
@ 34S of ore pyrite in the Getchell trend is near 0%, whereas in
the Jerritt Canyon district it is near 10% (Cline et al., 2005).
These relationships suggest that the source of the H2S in
ore fluids varied significantly within and between mining
districts.

(e)

100 000 13.15.3.5 Composition of Ore Fluids

Studies of fluid inclusions in ore-stage quartz from Nevada CTD
10 000 commonly indicate homogenization temperatures of 180–
240  C, salinities of 3–6 wt% NaCl (equiv.), and CO2 contents
of 1–4 mol%. After correction for the amount of CO2 present (see
succeeding text), salinities may actually be less than 3 or 4 wt%
1000
NaCl (equiv.). Ion chromatographic analyses of inclusion fluid
extracted from orpiment, realgar, and stibnite generally show that
Na > Cl > K> Mg > NH4 > F > Br; Ca and CO3 are compromised
100 by calcite contamination. Mass spectrometric analyses of
individual fluid inclusions or inclusion assemblages in ore-
stage minerals, opened by thermal decrepitation, show that
10 CO2 > CH4 N2 > H2S > SO2. Of particular importance for
(f) As Sb Au Cu Tl Ag Hg Pb estimation of ore-forming solution pH and gold transport,
CO2 is between 1 and 4 mol%, and H2S concentrations
Figure 21 (a–d) Photographs from Emsbo et al. (2003) of (a) barren are commonly 102–101 molal (Hofstra and Cline, 2000).
hydrothermal zebra texture dolomite in Devonian Bootstrap Limestone;
Eocene reconstructions of Carlin-type districts suggest most
(b) sulfidized dolomite with dark, fine-grained ore pyrite replacing ferroan
dolomite; and (c) internal sediment in a breccia cavity with layers containing
deposits formed at depths of 1–3 km, with a few as deep as
ore pyrite. The white line marks a laser ablation (LA) trace across one layer 5 km or as shallow as 300 m (Cline et al., 2005; and references
shown in (e). (d) Scanning electron microscope (SEM) backscatter image therein). Such depths correspond to hydrostatic pressures of
of euhedral pre-ore and bright anhedral ore pyrite. (e) Profile of arsenic less than 300 bars.
(red) and gold (yellow) concentrations in pyrite along the 8 mm LA
induction-coupled plasma (ICP)-mass spectrometer (MS) trace. The dips in
As and Au correspond to the position of coarse-grained pre-ore pyrite
grains. (f) Bar and whisker plot for nine elements detected in ore pyrite 13.15.3.6 Carbonate Dissolution
along the trace. The 10th, 25th, 50th, 75th, and 90th percentile values are
Carbonic acid appears to have been the major acid volatile in
indicated.
ore fluids responsible for carbonate dissolution and hydrolysis
of feldspars to illite and kaolinite. Figure 23 shows that disso-
lution of carbonate rocks generates CO2 such that fairly high
exhibit nearly identical enrichments (Figure 20(b) and 20(c)) confining pressures (e.g., 250 bars) are required to prevent
and likewise must have formed from H2S-rich fluids at condi- fluid immiscibility (boiling), for which there generally is little
tions similar to those of CTD in Nevada. evidence. It also shows that high fluid–rock ratios (>550–775)
 Geochemistry of Hydrothermal Gold Deposits 405

0.1 0.1
Ore Ore Pyrite
Pyrite

0.01 Detection limit 0.01

Artifact
Au/As (ppm)

Au/As (ppm)
E-3 0.001

Au opt
>1
1E-4 0.1−1 1E-4
0.01−0.1
0.001−0.01
<0.001
1E-5 1E-5
1E-4 1E-3 0.01 0.1 1 −20 −15 −10 −5 0 5 10 15 20
(a) As/S (ppm) (b) δ34SCDT ‰

Figure 22 (a) Plot of whole rock Au/As and As/S ratios converge towards that of ore pyrite with increasing gold grade; (b) Plot of whole-rock d34S
values and Au/As ratios converge towards that of ore pyrite from Hofstra et al. (2011). From the ‘Screamer’ sector of Betze-Post deposit, Carlin Trend.

are required to produce the central zones of calcite and dolo- 100–1000 ppb range over a wide range of near neutral to acidic
mite absent argillic alteration. Any calcite precipitated in open pH. Such high gold contents have actually been detected in
fractures from such neutralized ore fluids contains a mixture of quartz-hosted fluid inclusions from CTD in the Dian-Qian-Gui
hydrothermal fluid and country rock carbonate, which can area of China (Su et al., 2009a, 2009b). To attain such high
mask the stable isotopic composition of ore fluids. gold solubilities, the activity of H2S must be greater than that
along the pyrite–hematite or pyrite–magnetite buffers
(Figure 25(b)). The paucity of hematite and magnetite in the
suboxic to euxinic seafloor depositional environment of car-
13.15.3.7 Ore Fluid Composition and Precipitation
bonaceous pyritic sedimentary rocks may be one reason that
Mechanisms
ore fluids evolved to such high sulfidation states. Perhaps
The sulfidation state of the ore fluids (log fS2 15 to 10) is another reason is that H2S in early stage ore fluids replaces
constrained by the homogenization temperatures of fluid in- any hematite, magnetite, or other iron-bearing minerals in the
clusions and the common occurrence of orpiment and realgar host rocks with pyrite, causing successive aliquots of ore fluid
in the ore (Figure 24). The oxidation state of the ore fluids to be more H2S rich.
(log fO2 41 to 34) is constrained by the absence of hydro- Quartz, stibnite, realgar, orpiment, gold, and cinnabar all
thermal hematite, ubiquitous presence of pyrite, and the pre- have prograde solubility and precipitate from ore fluids during
dominance of CO2 over CH4 in fluid inclusions (Figure 25(a)). cooling (Hofstra, 1994; Hofstra et al., 1991). In contrast, cal-
The pH of ore fluids was moderately low as indicated by the cite has retrograde solubility and dissolves as ore fluids cool or
common occurrence of kaolinite or dickite, with or without are neutralized by carbonate rocks. Calcite may precipitate
sericite, and the general absence of adularia (Figure 25(a) and as neutralized ore fluids lose CO2 to the vapor phase. Barite
25(b)). The dissolution of calcite and dolomite is further evi- forms as Ba in reduced ore fluids and mixes with external fluids
dence that the early fluids were acidic with pH less than about containing marine sulfate derived from sedimentary rocks
4.5 (Figure 23). Marcasite, which occurs in some ores, is only (Cline et al., 2005).
stable at pH less than 5 and temperatures less than 240  C Equilibrium chemical modeling of isothermal reactions
(Murowchick, 1992). The low pH of ore fluids and consequent between H2S-rich ore fluids and host rocks containing different
increases in Al solubility explain the abundance of Al in dark or amounts of reactive iron rationalize several key features of the
weakly cathodoluminescent bands of drusy quartz (Rusk et al., ores (Figure 21). First, the gold grade and extent of minerali-
2008). The alternation of Al-rich and Al-poor bands in drusy zation vary as a function of the amount of reactive iron in the
quartz with constant d18O values may reflect changes in pH host rocks (Figure 26), with large, low-grade ore bodies in host
related to the successive influx and neutralization of acidic rocks with small amounts of reactive Fe (e.g., 0.25 wt%) and
fluids via carbonate dissolution. small, high-grade ore bodies in host rocks with large amounts
The mineral evidence of a high sulfidation state (Figure 24) of reactive Fe (e.g., 5 wt%). Second, regardless of the iron con-
is consistent with the measured H2S content of fluid inclusions tent of the host rocks, the gold/pyrite ratio is relatively constant,
(avg. 102 molal; Hofstra and Cline, 2000). Figure 25 shows corresponding to about 1000–1500 ppm Au in pyrite. Such
that at the indicated range of H2S activity and fO2 that concentrations are similar to the average abundance of Au in
gold solubility as Au(HS)21– and AuHSo complexes are in the ore pyrite from several Nevada CTD (Reich et al., 2005).
406 Geochemistry of Hydrothermal Gold Deposits

7
2.5 Calcite
0 CO2
6.5
2 De
cal
−.5 cifi
cat
Dolomite ion
1.5 6

1 −1 De H2O
do
Gases (log fugacity)

Minerals (log kg) lom 5.5

itiz
.5 ati
on
−1.5

pH
5
0
−2 pH 4.5
−.5

−1 −2.5
4

−1.5
−3 3.5

−2
−3.5 3
0 100 200 300 400 500 600 700 800 900
Fluid reacted (kg)

60 80 100 120 140 160 180 200

Temperature (⬚C)

Figure 23 Plot of chemical reaction between a 200 C ore fluid containing 4 mol% CO2 and 6 wt% NaCl (equiv.) with a typical carbonate rock
comprised of calcite (1 kg) and dolomite (0.2 kg) containing a dilute, pH 7, 60  C pore fluid (e.g. local ground water at 2 km depth). The model shows
how the abundance of calcite and dolomite, pH, and the fugacity of H2O and CO2 vary during fluid–rock reaction. During decalcification, pH remains near
4.5. During dedolomitization, pH drops to 4.3. Following dolomite dissolution, pH drops to about 3.5 (i.e., central zone of carbonate absent argillic
alteration). About 525 kg of ore fluid is required to dissolve 1 kg of calcite, and an additional 250 kg is required to dissolve the remaining 0.2 kg of
dolomite from Cline et al., (2005).

13.15.3.8 Element Substitution in Pyrite et al., 2003). Thus, the rare formation of gold nanoparticles
in arsenian pyrite may be an indication of fluid mixing.
Despite the apparent correspondence between equilibrium
Most importantly, the microanalytical results indicate that
chemical models and the amount of Au and pyrite precipitated
most of the gold in CTD precipitated in solid solution in
by sulfidation in the ores, experimental studies show that the
arsenian pyrite (Kesler et al., 2011; Reich et al., 2005). The
solubility of Au in stoichiometric pyrite is only about 3 ppm
correspondence between the modeled proportion of gold and
(Tauson, 1999). In contrast, the solubility of Au in arsenopy-
pyrite produced by sulfidation of host-rock Fe and the abun-
rite is much greater at about 30 000 ppm (Fleet et al., 1997).
dance of Au actually present in solid solution in arsenian pyrite
These experiments suggest that entry of Au into the pyrite
suggest ore fluids could not have been highly undersaturated
lattice is facilitated by As. Microanalytical studies (electron
with respect to native gold. It also allows that the ore fluids
microprobe analyzer (EMPA), secondary ion mass spectrome-
would have been saturated with native gold if arsenic had not
try (SIMS), x-ray absorbance near edge structure (XANES) and
been present in sufficient amounts to form arsenian pyrite.
high resolution transmission electron microscopy (HRTEM))
Regardless of whether ore fluids were saturated with native
of arsenian pyrite from several CTD (Hough et al., 2011; Kesler
gold and pyrite or auriferous arsenian pyrite, the consumption
et al., 2011; Reich et al., 2005; and references therein) show
of H2S during sulfidation of host-rock Fe (external Fe-bearing
that as As1 substitutes into the S site, there is significant
fluids) was the driving mechanism for their precipitation. Fur-
substitution of Au1þ into pyrite (Figure 27). These studies
thermore, the assemblage of trace elements typically detected in
also demonstrate that arsenian pyrite with discrete Auo nano-
ore pyrite (As, Sb, Cu, Tl, Pb, Hg, and Ag; Figure 22) all form
particles (see Hough et al. 2011) is actually rare and only
sulfide complexes. The fixation of such sulfide-complexed trace
present in samples with Au/As molar ratios greater than
elements in ore pyrite may be an indication of the relatively low
about 0.02 (Figure 27). An exemplary sample with abundant
temperatures or ore formation and the rapid rate of pyrite
gold nanoparticles is from the Screamer sector of the Betze-
precipitation during sulfidation, which also is suggested by the
Post deposit (Kesler et al., 2011), where arsenian pyrite precip-
small grain size and anhedral morphology of ore pyrite
itated by sulfidation of an external Fe-bearing fluid (Kesler
(Figure 15).
 Geochemistry of Hydrothermal Gold Deposits 407

− −
225 ⬚C, 500 bars, a Cl = 1, a H2S = 10 2
ion −30
nsat

−
-5 nde (As,S)liquid

l2

CO2(aq)
HSO4 SO4

uC
ur co
Sulf

HCO3
A

Kaol.

Musc.
pb

Kaol.
Al.
1p

K-spar
Hematite

Musc.
-10 Cooling
im ent −35
Orp ite y
bn H2S
Sti imon
logƒS2

t
An

10
Pyrite

00
logƒO2
-15 lg ar
Rea rite AuHS+ 100 ppb

pp
Sulfidation
Py otit
e HS− S2
−40

b
CO2(aq)
rrh HCO −
Py

Au
nic CH4(aq) 3

se

(H
-20 Ar

S) 2
10 ppb Mag.
te
yri

-
p
no git
e −45
se
Ar e llin
Lo Pyrrhotite

b
pp
-25 (a) 1 ppb

10
100 150 200 250 300 350
Temperature (⬚C) 3 4 5 6 7 8 9 10
pH
Figure 24 Log fS2 versus T diagram showing stability fields of Fe
minerals (bold black line), As minerals (colored fields), Sb minerals (blue
line), and As–S liquid (dashed lines). The white dot corresponds to the 225 ⬚C, 500 bars, a Cl− = 1, pH = 5.1
modeled ore fluid and the black arrows shifts induced by cooling and −35
S

1 pp
0.1

10 0
10 p

100
100
sulfidation from Simon et al. (1999).
−36

ppb

0 pp

00 p
b

ppb
pb
Hem.
−37

b
Arsenian pyrite is present in other gold deposit types, in-

pb
cluding LS and HS epithermal deposits and orogenic gold de-
−38
posits, and can be classified into two types: As1 pyrite where
Pyrite
As substitutes into the S site under reducing conditions and −39
Asþ3 pyrite where As substitutes into the Fe site under more
log ƒO2

Sulfidation
oxidizing conditions (Deditius et al., 2008; Kesler et al., 2011). −40
Such pyrites plot in different domains on an As–Au plot
−41
(Figure 27). Pyrites from orogenic, Carlin-type, and LS epither- Magnetite
mal gold deposits mostly occur in the As1 pyrite field where −42
Auþ1 substitutes into the pyrite lattice. In contrast, pyrites from
HS epithermal deposits plot in the Asþ3 pyrite field and may −43
contain native gold (Figure 27). This difference indicates that
−44 Pyrrhotite
pyrite in orogenic, Carlin-type, and LS epithermal gold de- (b)
posits formed under more reducing conditions than pyrite in −45
HS epithermal gold deposits. Thus, despite the presence of −5 −4 −3 −2 −1
high sulfidation state As-minerals in CTD, the predominance Log a H2S(aq)
of As1 pyrite suggests they are geochemically more like oro-
Figure 25 (a) Log fO2 versus pH diagram showing solubility contours
genic and LS epithermal gold deposits and not simply Cu-
for gold (thin red lines) relative to stability fields of Fe minerals (colored
poor, sediment-hosted, analogues of HS epithermal gold
fields), Al silicates (vertical black lines), predominance fields of sulfur
deposits. species (black dashed lines) and carbon species (blue dashed lines), and
the model ore fluid (white dot). (b) Log fO2 versus log a H2S(aq) diagram
showing the solubility of gold (thin red lines) relative to the stability fields
13.15.3.9 Source(s) of Ore Fluid Components of Fe minerals (shaded fields) and liquid sulfur (S). The white dot
corresponds to the modeled ore fluid and the black arrow shifts induced
Water: H and O isotopic data from each district are usually
by sulfidation (adapted from Hofstra AH and Cline JS (2000).
suggestive of either two or, in a few districts, three sources of
water (Cline et al., 2005; Hofstra and Cline 2000; Hofstra et al.,
2005; Hu et al., 2002). Evidence for groundwater that plots West Qinling) is from main ore-stage kaolinite as well as
near the meteoric-water line is typically obtained from analyses inclusion fluids in main and later ore-stage minerals.
of late ore-stage minerals and inclusion fluids. Evidence for CO2 : As described earlier, CO2 is the principal acid volatile
exchanged meteoric water that plots well away from the in ore fluids, and CO2 metasomatism of marine carbonate
meteoric-water line is generally obtained on main ore-stage rocks liberates additional CO2 that may mask the isotopic
kaolinite, quartz, and inclusion fluids. Evidence for a more signature of the source. Consequently, the C and O isotopic
deeply sourced magmatic or metamorphic fluid detected in data on altered rock with recrystallized carbonate minerals and
some deposits (Deep Star) and districts (Getchell Trend and late ore-stage calcite often reflect the isotopic composition of
408 Geochemistry of Hydrothermal Gold Deposits

variably exchanged meteoric water and carbonate host rocks H2S: As shown in the preceding text, H2S is the main ligand
(Hofstra and Cline, 2000; Hofstra et al., 2005; Hu et al., 2002). for Au transport in CTD. If it were derived solely from a
A few districts (e.g., Getchell) exhibit data arrays that extend to magmatic source, ore pyrite from the heart of each ore zone
lower C and O values that may reflect a deeper magmatic or should have d34S values near 0% and late ore-stage orpiment,
metamorphic source. realgar, and stibnite should have complimentary values that
differ by their lower temperatures of formation and fraction-
ation factors. Whereas the Getchell Trend has some data that
fits this model, most districts yield much higher or lower d34S
225 ⬚C, 1 m Cl−, 10−2 m H2S, values that require a sedimentary source (Cline et al., 2005).
pH 5, logƒO2 -40 Some districts, such as the northern Carlin Trend, exhibit high
d34S values (10%) in the heart of large high-grade deposits
WT% IRON
1000–1500 ppm 5.0 (Meikle and Betze) and low d34S values (0%) in peripheral ore
100 5.0 Au in Pyrite
2.0 zones (Screamer). This may indicate that the isotopic compo-
sition of S in ore fluids evolved due to fluid–rock reactions

Au oz ton-1
Au ppm

2.0 1.0 (particularly considering the abundance of pre-ore pyrite in

1.0 sedimentary host rocks) or mixing with external H2S-bearing
0.5
10 0.5 fluids (Hofstra et al., 2011).
0.2
He: Helium isotope data of inclusion fluid extracted from
0.25
0.1 ore and late ore-stage minerals are suggestive of a mixture of
crustal and mantle-derived He with R/Ra of 0.33–1.41 in the
1 Carlin Trend and Jerritt Canyon district (Hofstra et al., 2010)
0 20 40 60 80 and R/Ra of 0.14–1.47 in the Getchell Trend (Cline et al.,
Grams rock added to 1 l of ore fluid 2005). The He isotope data resemble those of Eocene felsic
Figure 26 Plot showing the modeled gold grade and mass of ore intrusions with R/Ra of 0.11–0.76 that also contain mixtures
produced by sulfidation of rocks with different iron contents. In each of He from crustal and mantle sources (Hofstra et al., 2010).
reaction path, the gold/pyrite ratio corresponds to 1000–1500 ppm gold in With such data, it is difficult to discriminate between He from
pyrite (reproduced from Hofstra AH and Cline JS (2000). felsic magmas and He in convecting fluids that obtained

1e+01
Field of As3+ pyrite & Au0

1e+00

1e-01
5
10
4⫻
As content of arsenopyrite

+
Au (mol%)

1e-02 High sulfidation

C As
.02
=0 pe
1e-03 C Au n-ty
a rli
C

1e-04 Low sulfidation

ic
en

1e-05
og

1 ppm
Or

Field of As1- pyrite & Au1+

1e-06
1e-04 1e-03 1e-02 1e-01 1e-00 1e-01 1e-02

As (mole %)
Figure 27 Plot of EMPA and SIMS data showing mol% Au versus As analyses of pyrite from Carlin-type (yellow field), low sulfidation (LS) (green line),
high sulfidation (HS) (magenta line), and orogenic (orange line) gold deposits. The solubility limit for Au1þ in arsenian pyrite determined by Reich et al.
(2005) is indicated by the black diagonal line. Carlin-type pyrite that plots below the line contains As1 that substitutes into the S site and Au1þ that may
substitute into the Fe site, whereas the few data that plot above the line contain nanoparticles of native gold (Auo). Arsenian pyrite from orogenic and LS
gold deposits is similar to Carlin-type pyrite, whereas that from HS deposits is distinct with As3þ substitution into the Fe site. This difference indicates
that pyrite in orogenic, Carlin-type and LS gold deposits formed under more reducing conditions than pyrite in HS gold deposits (adapted from Kesler et
al., (2011) and data in Zacharias et al. (2004), and Large et al. (2009)).
 Geochemistry of Hydrothermal Gold Deposits 409

components from crustal and mantle sources. The greater deposits allows for much lower ore grades, including some
mantle He component in ore fluids permits that there were examples of large ore bodies grading between 0.45 and
mafic magmas at depth and deep fracture systems, such as the 1.0 g t1 Au (e.g. northern Sonora, Mexico; Morro do Ouro,
Getchell fault. Brazil).
Pb: While the abundance of Pb in ore pyrite and late ore-
stage sulfides is low, there is enough to analyze by conven- 13.15.4.1.1 General geologic setting and genetic model
tional methods on well-prepared separates. SIMS can measure The orogenic gold deposits are found in three different broad
206
Pb, 207Pb, and 208Pb, but the reference isotope 204Pb is temporal/geotectonic settings (Goldfarb et al., 2001). First,
generally below detection. Such data (Hofstra et al., 2010; ores hosted in Late Archean greenstone were mainly formed
Tosdal et al., 2003) suggest the Pb in ore pyrite in Nevada between 2750 and 2520 Ma in granitoid–greenstone terranes
CTD is derived from Neoproterozoic siliciclastic sedimentary worldwide. This group of deposits includes those of the Yilgarn
rocks that underlie the host Paleozoic carbonate rocks. In craton (Western Australia), Superior province (Canada), Slave
contrast, late ore-stage sulfides commonly contain Pb derived province (Canada), Tanzania craton (central Africa), Dharwar
from the Paleozoic host rocks. craton (India), Zimbabwe–Kaapvaal craton (southern Africa),
Sr, Nd, and Os: These isotopes tend to reflect the isotopic and São Francisco craton (Brazil). Second, ores dated between
composition of the marine carbonate, terrigenous detritus, or 2100 and 1730 Ma are hosted by Paleoproterozoic sedimen-
sulfides in the host rocks and thus far have been of limited tary and volcanic rocks deposited on the margins to the
value (Hofstra and Cline, 2000). granitoid–greenstone terranes. Banded iron formation and fer-
Summary: The isotopic data generally support models in ruginous chert are commonly favorable host lithologies in the
which meteoric water convects to sufficient depths and temper- Paleoproterozoic terranes. Examples include the ores of west-
atures to promote substantial isotopic exchange between the ern Africa, northern South America, and Homestake (USA).
fluids and sedimentary rocks from which S, Pb, and perhaps A final group of ores formed between c.650 and 50 Ma in
Au and As(?) were scavenged. Such isotopically exchanged fluids marine sedimentary rocks adjacent to active continental mar-
mixed with unexchanged groundwaters during the late ore stage gins. These include the Mesozoic to Tertiary gold deposits of the
as the systems collapsed. The deposits and districts with evi- North American Cordilleran orogen, Colombian Andes, eastern
dence for more deeply sourced water and metasedimentary Russia, and the Otago area of New Zealand, and the latest
S (e.g., Chinese CTD) permit an epizonal orogenic model, Neoproterozoic–Paleozoic gold deposits of the Tasman orogen,
while those with deeply sourced water and 0% @ 34S permit a Central Asia orogenic belt, Pan African orogen, and Eastern
magmatic-hydrothermal model (e.g., some in Nevada and Cordillera of the main Andean range. Although processes
perhaps in West Qinling CTD). Despite the clear evidence for that form orogenic gold deposits are continuing today, the lack
a deeply sourced fluid in some well-studied deposits (e.g., of significant ore systems younger than c.50 Ma reflects the
Getchell), there also are well-studied high-grade deposits that absence of required time for uplift and unroofing of ores formed
lack any evidence for such deeply sourced water or magmatic typically at midcrustal depths (see succeeding text). Most Phan-
sulfur (e.g., Meikle in the Carlin Trend). Consequently, at this erozoic orogenic gold deposits occur in the fore-arc terranes
writing, one single integrated model does not appear to account adjacent to a subduction-related continental arc (Figure 1);
for the variability of observed geochemical and mineralogical however, some may occur in deformed back-arc basins (Sukhoi
data; this leads to continual debate on the origin of these Log, Russia; Bendigo, Australia).
deposits. The temporal distribution of the orogenic gold deposits
correlates with periods of supercontinent growth, including
Kenorland, Columbia, Gondwana, Pangea, and ongoing Ama-
13.15.4 Orogenic Gold Deposits sia (Goldfarb et al., 2010). Juvenile crust added to the cratons
represents a favorable source reservoir for gold, related metals,
13.15.4.1 Introduction
and volatiles that are concentrated to form the eventual ore
Orogenic gold deposits are most readily defined as those in deposits. The thick, buoyant, and low-density subcontinental
deformed metamorphic rocks. As their name implies, orogenic lithospheric mantle, characteristic of Paleoproterozoic and
gold deposits represent regional fluid flow that is inherent to older cratonic blocks, has tended to prevent uplift and erosion
orogeny. The gold-bearing ore bodies typically form in host of many Precambrian terranes. The Mesoproterozoic to early
rocks between 20 and 200 m.y. after deposition of the rocks Neoproterozoic orogenic belts of Rodinia generally lack eco-
and relatively late in their deformation history. Historically, nomically significant gold deposits. This reflects uplift and
these were thought of as a relatively high-grade type of gold erosion of these older accretionary-type orogens, such that
deposits, with gold-bearing quartz veins mined underground almost all orogenic gold deposits that would have formed
at grades of at least 5 g t1 and commonly at >10 g t1. With between c.1730–650 Ma were lost to the geologic record once
rising metal prices, improved mining methodologies, and the presently exposed, mainly deep-crustal high-grade meta-
more efficient gold extraction techniques, the character of the morphic rocks reached the surface. In addition, the growth of
ores has dramatically changed during the past few decades. Rodinia subsequent to the breakup of Columbia was likely a
Many of these deposits are now mined by large open-pit oper- gradual reconfiguration, rather than two distinct events. Dur-
ations, recovering both high-grade veins and surrounding ing such a reconfiguration, the Cordilleran-type orogenic belts
lower-grade hydrothermally altered wall rock that had been were dominated by mature sediments shed from the adjacent
considered waste, or even just distal geochemical anomalies, in cratons, rather than by more volatile- and metal-rich accreted
the past. Thus, present-day open-pit mining of orogenic gold oceanic sediments. This too may have hindered development
410 Geochemistry of Hydrothermal Gold Deposits

of significant gold resources during the Mesoproterozoic to deposits hosted by high-grade metamorphic rocks (e.g.,
Neoproterozoic ‘Boring Billion’ year gap. Hemlo, Challenger, Big Bell) formed during greenschist facies
Structure plays the major role in localization of most oro- events on a prograde metamorphic path (Phillips and Powell,
genic gold provinces, with ores typically being formed relatively 2009). These may reflect a complex and less common P–T path
late within an evolving orogen. Deposits are near steeply dip- type such that perhaps multiple or overlapping metamorphic
ping transcrustal fault systems that are parallel to the strike of events affected the same region.
the orogen or the greenstone belt. The first-order structures are The genetic model most widely accepted for the orogenic
many hundreds of kilometers in length and several hundred gold deposits is that they are a product of metamorphic devo-
meters wide, often being defined by a series of closely spaced latilization (see summary in Phillips and Powell, 2010). Dur-
distinct faults that reflect a cumulatively broad zone of defor- ing greenschist to amphibolite facies metamorphic reactions
mation. Ores are interpreted to have been deposited between 3 involving most sedimentary and volcanic rock sequences,
and 20 km depth, and thus subclassifications of ores as epizo- mainly at temperatures of approximately 440–520  C (Elmer
nal (<6 km), mesozonal (6–12 km), and hypozonal (>12 km) et al., 2006, Phillips and Powell, 2010), perhaps 3–5 vol% of
have been proposed by Groves et al. (1998). Irrespective of the rock is converted to a liberated fluid (e.g., Fyfe et al., 1978);
depth, a complex kinematic history is commonly represented the fluid volumes mainly depend on abundances of hydrous
by earlier reverse motion and later strike-slip events. The shift to and carbonate minerals. The produced fluid rapidly generates
a more transpressional to transtensional stress regime, and its own elevated transitory permeability (e.g., Tenthorey and
related seismicity, correlates with the majority of the gold- Cox, 2003) and is focused by resulting fluid pressure gradients
forming events (e.g., Goldfarb et al., 1991a, 1991b). into nearby deep-seated faults. The prograde metamorphic
Whereas the transcrustal fault zones are the main fluid con- fluid moves upward along these major structural channelways
duits that focus fluid flow (Cox, 1999; Fyfe, 1987), gold ore during episodic seismic events (e.g., Sibson et al., 1988). Gold
bodies are situated along second- and third-order subsidiary and associated metals are carried with the fluid phase to shal-
faults. These lower-order faults are dominated by brittle–ductile lower deposition sites in rocks already on a retrograde meta-
to ductile shear zones, but less commonly gold-bearing exten- morphic path (e.g., Stuwe, 1998), commonly located in
sional veins and vein arrays may also develop. Rheologic hetero- greenschist facies brittle–ductile regimes.
geneities result in a variety of lithologies in greenstone belts and Gold is sourced and released from elevated concentrations
metasedimentary rock-dominant terranes serving as favorable ore in preexisting pyrite, which is syngenetic to diagenetic in origin
host rocks, particularly in areas of minimum mean stress, such as in sedimentary rocks (Goldfarb et al., 1997), during the pro-
fault jogs, fault bends, and fold hinges. Ore-bearing quartz– grade desulfidation to form pyrrhotite. In greenstone belts, the
carbonate veins are commonly laminated or contain breccia pyrite might be present in altered areas of metabasic rocks or
fragments, which reflect repeated fracturing and fracture sealing within minor, but widely distributed chert beds; for example,
in the hydrothermal systems (e.g., Cox et al., 2001) over periods near the Golden Mile deposit, Western Australia, the highest
of a few million years (e.g., Miller et al., 1994). Larger ore bodies gold background measurements are associated with seafloor-
have extensive down-plunge continuity of as much as 1–2.5 km. altered serpentinite (16–87 ppb Au) and interflow sedimentary
Geologists have long recognized that most orogenic ores rocks (16–420 ppb Au) (after Golding, 1978, as shown in
occur in greenschist facies rocks (e.g., Boyle, 1979; Buryak, Table 5 of Shackleton et al., 2003). The widely disseminated
1964; Groves et al., 1998). Barrovian metamorphic sequences, pyrite grains contain a few hundred ppb to a few thousand ppb
with inverted geothermal gradients reflecting the contraction- gold (Boyle, 1979; Thomas et al., 2011). Some workers have
related thrust faulting of high-grade rocks over lower-grade argued for the need for metamorphism of a gold-enriched
rocks, are inherently characterized by an abundance of aurifer- protolith to form large gold provinces, such as proposed buried
ous quartz carbonate veins in greenschist zones (Figure 28). ‘fertile’ volcanic sequences below the turbidite-hosted ores in
Where large veins exhibit high densities along major transcrustal the Victorian goldfields (e.g., Bierlein et al., 1998; Willman
structures, giant gold districts exist; where generally smaller vein et al., 2010), but it appears that adequate fluid and sulfur
systems are scattered throughout the greenschist facies, without budgets should ensure that such a model may characterize
any apparent relationship to the first-order faults, lode deposits most typical active margin lithologies and a specific favorable
are smaller and commonly subeconomic (e.g., Klondike placer Au-enriched source horizon is not required.
district; MacKenzie et al., 2008). Rocks with different metamor- Supporting evidence for the metamorphic model includes
phic grades may be exposed at different locations along the the following: (1) the common H2O–CO2–CH4–N2–H2S, low-
length of a major fault system, and gold is restricted to the salinity fluid that is typical of a product formed during devo-
greenschist environments. For example, in southeastern Alaska, latilization; (2) the isotopically heavy d18O and dD values for
deposits of the Juneau gold belt are hosted by greenschist facies hydrothermal minerals (Goldfarb et al., 1991a; Kerrich, 1989);
rocks along the northern quarter of a transcrustal fault system (3) the variable d34S values that reflect variable host terranes
(Goldfarb et al., 1988). Amphibolite and granulite facies rocks undergoing a regional metamorphism (e.g., Goldfarb et al.,
exposed along the same structure to the south are cut by quartz 1997); and (4) the general lack of ores in high-grade metamor-
veins, but these lack anomalous gold concentrations (Goldfarb phic rocks because at relatively high temperatures, most mo-
et al., 1997). bile sulfur and gold have already been removed from the rock
Almost all orogenic gold deposits formed subsequent to the sequence. A gradual decrease in ore-related metals at increasing
development of the metamorphic mineral assemblages within metamorphic grades in a Barrovian sequence (e.g., Pitcairn
the immediate country rocks and thus on the retrograde meta- et al., 2006; Figure 29) supports a model where the ore-
morphic curve. A few, particularly Precambrian, large gold forming materials are derived from the accreted crustal
 Geochemistry of Hydrothermal Gold Deposits 411

Explanation
A⬘ Qtz. Diorite
e
Zon Diorite/Gabbro
nite
Sillima Alkali Gabbro
Isograd
R.

ez one
na

ndin
sit

a Fault
Su

Alm
Lode Gold
Placer Gold
e
Zon
r.

ite
C

Biot
ez
ld
Va

0 1 2 3 km

Metamorphic Isograd Ag As Sb Bi B
one Prehnite–Pumpellyite zone 64.50 6.93 0.60 0.08 2.18
rite z
Chlo
Chlorite zone 90.25 8.33 0.40 0.09 6.50
Biotite zone 61.29 3.01 0.09 0.08 0.79
ne
te zo
Almandine zone 72.20 0.31 0.10 0.07 0.70
llyi Sillimanite zone 42.80 0.52 0.03 0.03 0.70
mpe
e–Pu
hnit
PreWindy Cr. Prehnite– Silliimanite
Chlorite Biotite Almantine
pumpellite zone
zone zone zone
zone
Elevation Valdez
Windy Cr.
3000⬘ Cr.

0⬘
A A⬘

Figure 28 Map and cross section showing the spatial relation of the Valdez Creek district in central Alaska, which contains a series of small orogenic
gold deposits and associated placers, to an inverted metamorphic sequence within a clastic sedimentary rock terrane (after Smith, 1981). Similar to
the majority of orogenic gold deposits, those in the Valdez Creek district are spatially associated with the greenschist facies rocks (Goldfarb et al., 1997).
The source for the gold and related metals is the sequence itself, with pyrite to pyrrhotite conversion at temperatures above middle greenschist,
releasing sulfur and metals into an evolving fluid. Geochemical analyses of trace elements in the country rocks indicate a gradual decline in As, B, Bi, Sb,
and, to a lesser degree, Ag, as metamorphic grade of the rocks increases; similar patterns would be expected for Au and W, but in this study most
analyses for these elements were below analytical detection limits (E. Marsh, unpub. data).

allochthons tens of millions of years to a couple of hundred of 1:1–10:1, which contrasts with more Ag-rich LS epithermal
million years after the rocks were initially deposited. However, type deposits. In addition to electrum, gold may be present
orogenic gold deposits in the North China craton and Sonora, in maldonite, calaverite, petzite, or aurostibite, reflecting en-
Mexico, which are hosted by high-grade metamorphic rocks richments of Bi, Te, or Sb in many hydrothermal systems. The
billions of years older than Cretaceous gold lodes, indicate in typical range of 1–20 ppm Au, which is representative of most
some circumstances that aqueous-carbonic fluid and ore com- economic ore bodies, indicates three orders of magnitude of
ponents must be derived from metamorphosing slabs or their enrichment of gold relative to the 1–5 ppb Au range of most
overlying sediments being subducted below reactivated craton crustal source rocks. If one assumes that juvenile or sedimen-
margins. This is consistent with models that indicate a signif- tary pyrite with background concentrations of a few hundred
icant volume of CO2 flux is derived from decarbonation of ppb to be the source of the gold in the ore deposits (see
subducting slabs (Gorman et al., 2006). preceding text), gold in the veins is enriched 10–20 times
that in source minerals themselves.
Gold-bearing quartz  carbonate veins typically contain
13.15.4.2 Geochemistry and Mineralogy of
2–5% sulfide minerals. These are dominantly pyrite in igneous
Alteration and Ores
rocks and arsenopyrite in metasedimentary rocks. Pyrrhotite
13.15.4.2.1 Ore mineral assemblages becomes more abundant in higher temperature deposits or in
The geochemistry of orogenic gold deposits is well defined. relatively reduced ore host rocks, including most banded ion
Gold, most consistently present in electrum, occurs in veins as formations (BIFs). Trace element enrichments are most consis-
grains within quartz, as grains in contact with country rock tently As, Ag, Sb, B, Hg, Te, Bi, and/or W (Figure 30). Although
laminations or within altered wall rock fragments, or as remo- arsenic is mainly in arsenopyrite, it also occurs less commonly
bilized veinlets cutting originally coeval sulfide minerals. In in loellingite in hypozonal ores and realgar and orpiment in
altered wall rock, gold may be chemically bound along the epizonal ores (e.g., Groves et al., 1998; see Figure 24). The
outer parts of arsenopyrite or arsenian pyrite grains. The fine- epizonal ores may also be characterized by stibnite and traces
ness of most ore bodies is relatively high, with gold:silver ratios of cinnabar, reflecting the solubility of Hg and Sb as sulfide
412 Geochemistry of Hydrothermal Gold Deposits

Schist Wakamanna Endeavour Depletion Enrichment

Hg inlet As Se Sb W Ag Au Mo Cd Hg Otago
Au Malborough
Hot springs
W pin
e schist (Hg)
Al ist 100
sch
Sb Nentharn

Temperature (⬚C)
200 (Au, Sb)
Barewood
CHRISTCHURCH (Au, Sb)
300 Bonanza
Glenorchy (Au, As)
Macraes
400 (Au, W, As)
Macraes flat
Bonanaza 500
Otago
schist Nanthorn
Barewood 600
Waipori 0 200 km
Waitaka
Waitahuna

Figure 29 (Left) Map showing location of the Haast schist of the South Island, New Zealand, and that host orogenic gold and associated deposits.
(Right) Plot showing elemental enrichments and depletions in various gold deposits within the metasedimentary rocks at metamorphic grades reflecting
temperatures higher than gold ore formation (both diagrams are reproduced from Pitcairn IK, Teagle DAH, Craw D, Olivo GR, Kerrich R, and Brewer TS
(2006) Source of metals and fluids in orogenic gold deposits. Insights from the Otago and Alpine schists, New Zealand. Economic Geology 101:
1525–1546).

complexes in hydrothermal fluids at relatively low tempera-

Si Amphibolite facies Si
tures and thus their selective transport into the upper few Greenschist facies
kilometers of the crust along the shallower levels of the re- Ag Ag
gional fault systems. The presence of tourmaline in many veins K, Rb, Ba Na
Felsic rocks
Depletion
is consistent with elevated B levels in many ore bodies. Tung-
Na− Bi, Se
sten is most consistently in scheelite, perhaps concentrated by

Unaltered
Unaltered

CO2 complexing. Tellurides may be the dominant ore-hosting Se Bi, Se CO2

phase in Precambrian (e.g., Golden Mile, Western Australia) or Ba; Y-depletion Au
Phanerozoic (e.g., Kensington, Alaska) orogenic gold deposits.
Bismuth may be present in solid solution in galena or in traces Au, K, Rb K, Rb, Ba
of Bi-rich mineral species (e.g., bismuthinite and maldonite). CO2− W
The majority of the deposits are characterized by Cu and/or W− As, Sb
Pb–Zn concentrations in the hundreds of ppm range, reflecting
As, Sb− Te
trace amounts of chalcopyrite, galena, sphalerite, or tetrahedrite.
Te−

Intermed
Limited base metals and a common Ag–As–Au–Sb–Hg geo-
chemical signature reflect low-salinity ore-forming fluids (e.g., Schematic, not to scale Distal

Prox
ORE
Kerrich, 1983) and auriferous hydrothermal systems dominated Pasi Eilu 2001
by sulfide complexing (see succeeding text). Zoning of ore-
related minerals or gold grades is uncommon down-dip or Figure 30 Summary diagram from Eilu and Groves (2001) for
along strike in these deposits, reflecting limited temperature geochemical anomalies in ore bodies and surrounding mafic volcanic
variation (e.g., Kerrich, 1987; McCuaig and Kerrich, 1998). country rocks of the Yilgarn craton, western Australia. Most orogenic
gold deposits are enriched to some degree in Ag, As, Au, B, Bi, Hg, Sb,
13.15.4.2.2 Wall rock alteration Te, and W. Broader alteration halos for some trace elements relative to
Visible wall rock alteration is well developed in many relatively gold suggest they may be better pathfinders in the exploration for
reactive country rock types, but ore-hosting turbidite sequences orogenic gold deposits.
typically show only weak alteration. The most extensive alter-
ation zones are located where fluids reacted with granitoids alteration halos. Modern mining methods now consistently
and serpentinites in terranes dominated by metapelites or develop such auriferous-altered wall rocks as low-grade parts
greenstones. Sulfidation, (de)-silicification, carbonization, to the main ore bodies. Sulfidation and gold deposition in wall
and sericitization are the most consistent types of alteration, rocks reflect the sulfide/carbonate ratio of the hydrothermal
reflecting mobility of S, Si, K, and CO2 within the hydrother- fluid and the Fe/(Fe þ Mg) ratios of the wall rocks (Böhlke,
mal systems. 1988). Silicification of wall rocks is represented by branching
Pyrite and arsenopyrite are the most common wall rock veinlets of quartz trending outward from the main ore zones.
sulfide phases, with abundance decreasing away from ore Silica depletion halos in country rocks surrounding most ore
zones. Sulfide precipitation in wall rocks destabilizes bisulfide zones indicate some of the vein silica is locally derived. Sericite
complexes, causing deposition of wall rock gold in the to coarser muscovite characterizes many alteration
 Geochemistry of Hydrothermal Gold Deposits 413

assemblages, with visible bleaching as mafic minerals are Salinities are usually low, in the range of 1–3 wt% NaCl
replaced. In more mafic rocks, chlorite may be dominant at (equiv.). There are some deposits where observed ore fluids
low temperatures and biotite at higher temperatures. Within are solely carbonic (e.g., Ashanti, West Africa, in Schmidt-
ultramafic host rocks, fuchsite can be the most widespread Mumm et al., 1997; Carara, Guiana Shield, Brazil, in Klein
mica phase. Slight increases in ratios of alkalis to hydrogen and Fuzikawa, 2010). These are probably rare cases where
ion in the hydrothermal fluid can lead to significant amounts fluid immiscibility has taken place somewhere along the hy-
of K-feldspar or albite deposition. Hydrothermal carbonates drothermal flow path and gold has been transported to the site
include ankerite, siderite, calcite, dolomite, and magnesite, of deposition within the separate, more volatile unmixed
with compositions reflecting the bulk compositions of the component.
rocks. In some districts, an early district-wide carbonization Gold solubility is clearly as a complex with reduced sulfur
event has been reported to proceed the main period of ore (e.g., Loucks and Mavrogenes, 1999; Mikucki, 1998; Seward,
formation. 1973). The main ligand is HS and the dominant species in
There are some hypozonal deposits, particularly in the solution is AuHSo in slightly acidic solutions (e.g., Figures 10
Yilgarn, Dharwar, and Superior cratons, where alteration min- and 25). At typical hydrothermal temperatures of 250–400  C,
erals include biotite, Ca-amphiboles, plagioclase, pyroxene, AuHS becomes the more significant species in neutral to
and/or garnet. These reflect stable secondary hydrothermal slightly alkaline fluids. Most estimates indicate neutral to slightly
silicate minerals deposited surrounding ore bodies formed at alkaline solutions of pH 5–6 (e.g., Reed, 1997) such that both
higher temperatures, which is explained in detail by the ‘holis- gold species are probably important to varying degrees. In these
tic continuum model’ concept of Groves (1993). Also at rela- relatively low-salinity fluids that typify orogenic gold deposits,
tively high temperatures and X(CO2), relatively oxidized AuCl2 is not an important species except in rare higher temper-
mineral phases, such as hematite, magnetite, and anhydrite, ature hydrothermal systems. Total sulfur concentrations may be
may be stable (Evans et al., 2010). as high as 103.5 molal for epizonal deposits or 1 molal for
hypozonal ores (Mikucki, 1998). The gold itself is dissolved in
13.15.4.2.3 Geochemistry of ore-forming fluids the H2S-bearing fluid as the fluid is being formed via devo-
The chemistry of the hydrothermal fluids responsible for for- latilization (Phillips and Powell, 2010), and pyrite is being con-
mation of orogenic gold deposits is unique when compared to verted to pyrrhotite (e.g., Ferry, 1981; Tomkins, 2010). About
those that form other hypogene ores. The fluid, although H2O- 3–15 t Au km3 of rock may be dissolved assuming 80% effi-
dominant, contains 5–25 mol% nonaqueous species domi- ciency of gold extraction (e.g., Phillips et al., 1987). The elevated
nated by CO2 (Figures 31 and 32). The CO2 contents tend to CO2 in these fluids forms a weak (carbonic) acid and may help
be near the low end of this range in Phanerozoic systems and buffer the ore fluid in a pH range where gold is most soluble in
the high range in Archean deposits, which could reflect slightly sulfide complexes (e.g., Phillips and Evans, 2004).
higher Archean temperatures for the ore fluids and/or greater Oxygen isotopes (d18O), most commonly measured on ore-
relative CO2 production via devolatilization from greenstone hosted quartz, range between about þ10 and þ17%, and
belts. In more reduced hydrothermal systems, methane and calculated fluids are typically þ5 to þ10% (Figure 33). As a
nitrogen are present at percent levels and, in some deposits, generalization, the slightly higher fluid values in the range are
dominate over CO2. Whether this reflects differences in source more characteristic of Phanerozoic deposits, reflecting either
reservoirs or redox reactions at deposition sites is unclear. lower ore-forming temperatures and/or the metasedimentary

NaCl

0.6

Potassic
Weight 0.4 Porphyry
fraction High-sulfidation
NaCl Epithermal
Phyllic Low-sulfidation
0.2 Lode-gold

H2O CO2
0.2 XCO 0.4
2

Figure 31 Plot showing general ranges of compositions of orogenic gold deposits and some other deposit types. Note that orogenic ores are
characterized by a low to moderate salinity, CO2-rich ore-forming fluid. Most commonly the nonaqueous volatile content is about 5–25 mol%. These
contrast with fluids that are from epithermal and porphyry deposits that lack significant concentrations of nonaqueous volatiles and can be much more
saline (reproduced from Ridley JR and Diamond LW (2000) Fluid chemistry of orogenic lode gold deposits and implications for genetic models. Reviews
in Economic Geology 13: 141–162).
414 Geochemistry of Hydrothermal Gold Deposits

rather than metavolcanic rock source region. Hydrothermal isotope fractionation between the hydrothermal fluid and
flow along major fault zones is highly fluid dominated, but structurally bound water adds further uncertainty to dD fluid
some exchange obviously takes place at sites of ore deposition. inclusion measurements (Simon, 2001).
This is supported by the fact that in many districts where some Sulfur isotopes are quite variable between deposits/districts
veins are hosted in felsic intrusions, measurements of d18O are of different ages and tend to reflect the sulfur signatures of the
a few per millimeter lighter than those for auriferous quartz rocks undergoing devolatilization (e.g., Chang et al., 2008;
within other lithologies. Goldfarb et al., 1997; Figure 34). In individual large deposits,
Hydrogen isotopes, when measured for hydrothermal significant variation between ore bodies perhaps reflects
micas, consistently indicate fluid dD of 20 to 35% multiple mechanisms of gold precipitation (Hodkiewicz
(Figure 33). Most studies using dD values from bulk extraction et al., 2009).
of fluid inclusion waters yield much lighter ratios, but because Carbon isotopes for ore fluids have been determined from
of the typical multiple generations of fluid inclusions trapped many deposits of all ages from measurements on hydrothermal
in the ore bodies during tens of millions of years of postforma- carbonates. Most estimates for d34C are between 0 and 10%
tional deformation and uplift, these measurements are mean- and are not diagnostic of a specific source; these data can reflect
ingless (e.g., Pickthorn et al., 1987). In addition, hydrogen both mantle and many crustal carbon reservoirs.

NaCl NaCl

0.2 Mixed low-salinity

aqueous-carbonic
Low-salinity
aqueous Carbonic

H2O CO2(+CH4)
0.2 0.4 0.6 0.8
93%

Recorded at deposits
- by percentage

27% 38%

Figure 32 Compositional plot showing that most orogenic gold deposits form from low-salinity, mixed aqueous-carbonic fluids. Fluid immiscibility is
recorded in about one-third of the deposits studied by fluid inclusion observations (Ridley and Diamond, 2000). More commonly, fluid immiscibility
characterizes Precambrian rather than Phanerozoic deposits, likely reflecting slightly higher ore-forming temperatures or slightly more CO2-rich ore
fluids developed in earlier greenstone belts relative to younger accreted metasedimentary rock-dominant terranes.

Evolution paths for

tertiary meteoric waters
0 X = interior alaska
= southern alaska
Orogenic
Au deposits
Primary
0.01
-50 in
e 0.01
rl
Magmatic
d D (‰)

te H2O
wa
ric 10 1 0.1
eo
et
-100 M
Nevada sediment-hosted
Nevada
Au deposits 0.1
epithermal
deposits
10 1
-150
-20 -15 -10 -5 0 5 10 15

d 18O (‰)

Figure 33 Plot of dD versus d18O of ore-forming fluids showing that orogenic gold deposits are formed from a unique fluid that is isotopically heavy
in both oxygen and hydrogen as compared to epithermal and Carlin-type deposits. Most fluids are characterized by oxygen values of 5–10% and
hydrogen values between 20 and 35%. These are much heavier in hydrogen than meteoric or mixed meteoric and deep-crustal waters that form
Carlin and epithermal gold deposits.
 Geochemistry of Hydrothermal Gold Deposits 415

Measurements of d15N on hydrothermal micas cover a wide Some other volatile species, such as boron and noble gases,
positive range and have been interpreted as consistent with have been measured within orogenic gold deposits to deter-
metamorphic devolatilization processes (Jia and Kerrich, mine fluid sources, but the meaning of the resulting data are
1999; Jia et al., 2003). unclear. Boron isotopes in gold-related tourmaline in an Ar-
Heavy radiogenic isotopes such as lead have proven to be chean greenstone belt in India suggest boron was derived from
less diagnostic of orogenic gold ore-forming processes than both metamorphic devolatilization and I-type magmatism
the stable isotopes. Many workers have tried to use such (Krienitz et al., 2008), but does not necessarily imply that
isotopes to suggest source rocks for the gold ores, assuming gold and other volatiles have the same multiple-source charac-
elements such as lead, strontium, and neodymium have fol- teristics. Noble gas data often show a mantle contribution but
lowed the same evolutionary and hydrothermal path. Lead even in deposits where, in contrast, sulfur was clearly derived
isotope data for gold deposits from the Alaskan Cordillera from the crust (e.g., Dongping, northern China, in Mao et al.,
(Goldfarb et al., 1997; Figure 35) and from the Eastern 2003). Because many orogenic gold deposits are associated
Cordillera of Peru (Haeberlin et al., 2003, 2004) indicate with first-order structures that reach down to the mantle, the
signatures that reflect a strong lead contribution from sink presence of noble gases from the mantle, as well as lampro-
(i.e., variable host-rock lithologies) areas, rather than from phyric magmas, within these conduits is not unexpected. How-
metal source areas. Mortensen et al. (2010), however, noted ever, this does not suggest that the main ore fluid components
two distinct radiogenic Pb clusters for sulfides within aurifer- and the metals also have mantle sources. Fairmaid et al. (2011)
ous veins of one general rock type, the Otago Schist on South used noble gas and halogen data from Ballarat, Victoria, to suggest
Island, New Zealand. They interpreted these data to indicate signatures that were mixtures of deeper devolatilization-related
mobilization of at least minor amounts of lead by the hydro- fluids in a main source area and a more local contribution from
thermal fluids, although both clusters were interpreted to near the site of ore deposition.
reflect metals that were derived from the upper crustal schist
belt and the isotopic differences being reflective of distinct 13.15.4.2.4 PTX constraints on ore deposition
earliest and latest Early Cretaceous ore-forming events. In the Orogenic gold deposits reflect hydrothermal systems that pre-
Precambrian greenstone belt, lead isotopes may better finger- cipitate gold over a wide range of temperatures and pressures.
print metal source reservoirs because of the low lead concen- This is a consequence of the fact that these gold ores are
trations in the greenstone sink areas. In the Yilgarn craton, products of regional geothermal gradients and not local mag-
lead isotopes for orogenic gold deposits suggest lead, and matic heat engines, such as those that control most epithermal
thus perhaps gold, was sourced from basement granite–gneiss type gold deposits. Published ore formation temperature esti-
terranes (Qiu and McNaughton, 1999). The usefulness of mates are spread between about 225 and 600  C, although
strontium as a metal source region tracer also appears to most consistently fall in the 275–350  C range, and pressures
vary from one study to another (see discussion in Goldfarb range from 1 to 5 kb. Continuous fluctuation between litho-
et al., 2005). static and hydrostatic pressure regimes during flow migration

0
7
6 Oceanic
sulfate
200 5
4

400
age (Ma)
Host

1 2 3
1 Nome
2 Fairbanks district
600 3 Mohawk mine,
Fairbanks district
4 Juneau sed-hosted
5 Juneau igneous-hosted
800
6 Valdez Creek and
Windham Bay
7 Chugach terrane
1000
−30 −20 −10 0 10 20 30
d 35S (‰)

Figure 34 Plot of d34S values for sulfide minerals from orogenic gold deposits versus age of host-rock terranes for major gold districts in Alaska
(reproduced from Goldfarb RJ, Miller LD, Leach DL, and Snee LW (1997) Gold deposits in metamorphic rocks of Alaska. Economic Geology Monograph
9: 151–190). Oceanic pyrite in these terranes would reflect the oceanic sulfate curve for the past billion years, but with values about 20% lighter. The
sulfur isotopes of the minerals in the deposits reflect the signature of sulfur in the terranes such that heavy sulfur in sulfides reflects deposits formed in
terranes consisting of rocks deposited when seawater sulfate was heavy and vice versa. This is consistent with the metamorphic devolatilization model
for orogenic gold.
416 Geochemistry of Hydrothermal Gold Deposits

15.70
Fairbanks Circle Circle
Chugach-Kenai Mtns 39.2
Nome Juneau
Brooks Fairbanks
15.66 (Alaska-Juneau)
Range
Brooks Circle
Range Nome
Juneau
207Pb/204Pb

(Alaska-Juneau) 38.8
15.62
Circle

208Pb/204Pb
Chugach Willow Creek Willow Creek
(McHugh) Chugach-
Chugach Chugach
Kenai Mtns
15.58 (granite-related) (granite-related)
Willow Creek
(granite-hosted) 38.4 Chugach
Willow Creek
(granite-hosted) (McHugh)

15.54 Juneau
(Treadwell) Juneau
(Treadwellz)
38
18.4 18.8 19.2 19.6
18.4 18.8 19.2 19.6
206Pb/204Pb
206Pb/204Pb

Figure 35 Plot of lead isotope analyses for sulfide minerals from orogenic gold deposits in Alaska. Within the same gold province (e.g., Alaska Juneau
and Treadwell in Juneau gold belt; McHugh melange, S-type granite, and Chugach-Kenai Mountains metasediments of Chugach terrane), Pb isotope
measurement vary with changing host-rock lithology. This indicates that much of the lead is being contributed at the site of ore formation and says little
about the source of gold that has been transported to the deposit area by the hydrothermal fluid (reproduced from Goldfarb RJ, Miller LD, Leach DL, and
Snee LW (1997) Gold deposits in metamorphic rocks of Alaska. Economic Geology Monograph 9: 151–190).

along fault zones is responsible for measurements of huge Fluid–wall rock interaction is a more common gold depos-
pressure fluctuations at sites of gold deposition. Where hydro- iting process in the Archean, reflecting sulfidation of Fe-rich
thermal fluids reach levels shallower than about 3 km and thus wall rocks in many greenstone belts (Lambeck et al., 2011) and
temperatures less than about 225  C, progressive water–wall abundant sulfide mineral and high-grade gold precipitation
rock interactions raise pH so that gold is no longer very soluble adjacent to quartz veins. Oxidation of the fluid during sulfida-
in reduced sulfur complexes; Hg- and Sb-rich sulfides predom- tion may cause a drastic drop in gold solubility. In many
inate over Au in such epizonal environments (e.g., Goldfarb Phanerozoic and Paleoproterozoic gold provinces, there is an
et al., 2004), and the mercury may be carried to very shallow association of ores with carbonaceous and organic material.
levels in a vapor phase (e.g., Morteani et al., 2010). Ore for- Gold deposition in these environments may be related to
mation has a maximum permissive temperature of about reduction of the ore fluid by the carbon (e.g., Gize, 1999).
600  C, which is where fluids begin to become involved with Shifts in pH and temperature decreases are not important
partial melting (e.g., Tomkins and Grundy, 2009). factors in the formation of orogenic gold (e.g., Phillips and
The main controls on ore deposition are pressure fluctua- Powell, 2010). Fluid mixing has been invoked by some
tions and water–rock reaction. Drops in pressure are particu- workers (e.g., Neumayr et al., 2005, 2008) as important for
larly significant in metasedimentary rock-hosted ore systems, orogenic gold formation based on observations of oxidized
where the highest ore grades are within the vein systems them- and reduced mineral phases in deposit halos. However, car-
selves. Loucks and Mavrogenes (1999) noted a 2 kb drop bonation and sulfidation of ferric iron-rich minerals, such as
in pressure during hydrofracturing at 400  C and 10 km magnetite, can produce alteration assemblages with pyrite in
depth will cause a 90% decrease in gold solubility. Such pres- equilibrium with hematite and sulfate (Evans, 2010; Evans
sure drops may trigger CO2 immiscibility (e.g., Naden and et al., 2006).
Shepherd, 1989), also increasing pH, which further destabi-
lizes gold-bearing complexes. The common association of gold
13.15.4.3 Geochemistry of Type Examples
and tellurides in many orogenic gold deposits could be the
result of Te also partitioning into the nonaqueous phase during 13.15.4.3.1 Phanerozoic metasedimentary-hosted Muruntau
immiscibility in the form of H2Te (e.g., Tombros et al., 2010). The Muruntau deposit, located in the eastern part of the Tien
However, in many fluid inclusion studies of Phanerozoic Shan of Uzbekistan, is the largest known orogenic gold
deposits, there is a lack of clear evidence of fluid immiscibility, deposit. Discovered in 1958, it has a total production and
reflecting a nonaqueous component of perhaps only 4–6 mol% resource tonnage of about 5300 t Au at a grade of 3.5–4 g t1.
and thus a fluid remaining in the one-phase field for relevant It has been developed by open-pit methods since 1967. The
salinity and P–T solvi. ores are hosted by the Early Ordovician ‘variegated’ Besopan
 Geochemistry of Hydrothermal Gold Deposits 417

carbonaceous phyllite, with numerous Late Carboniferous to whereas Ne, Kr, and Xe in the sampled fluid inclusions are of
Early Permian granitic stocks scattered throughout the region. atmospheric origin (Graupner et al., 2006). Bulk extraction
Deposits mainly consist of relatively flat-lying quartz veinlets fluid inclusion @ 13C carbon isotope values between 5.3 and
and stockworks, although there are also relatively steep central 2.1% and log Br/Cl between 3.23 and 2.64 are both inter-
veins (Berger et al., 1994; Wilde et al., 2001). A Re–Os date preted as suggestive of a juvenile contribution to the hydrother-
on ore-stage arsenopyrite indicates mineralization at mal system (Graupner et al., 2006). There are, however, as
287.5  1.7 Ma (Morelli et al., 2007) and overlaps with a described earlier, always concerns on the significance of the
279  18 Ma Sm–Nd isochron age for hydrothermal scheelite analyses of bulk extraction fluid inclusion waters and noble
(Kempe et al., 2001). There are many possibilities as to why gases that may have different source reservoirs than other com-
Muruntau may be so large. These may include the following: ponents in a hydrothermal system. The d13C values for carbo-
(1) the occurrence of an exceptionally prominent jog in the main naceous matter in the unmineralized host rocks are 28.8 to
deep-crustal fault system, which separates the South and Central 28.0%, whereas in the ore bodies, the same material has only
Tien Shan just to the north of Muruntau; (2) an inherent high been enriched to values of 27.9 to 24.8% (Rusinova and
background of syngenetic/diagenetic gold in the local stratigra- Rusinov, 2003).
phy, as stressed by various Russian workers (e.g., Dolzhenko,
1991; 280 ppb estimate in Yermolayev et al., 1995); and/or (3) a 13.15.4.3.2 Paleoproterozoic BIF-hosted Homestake
thick sequence of carbonate rocks in the local geology, which The Homestake deposit in the Black Hills, SD, USA, is the
may have provided a pathway for hydrothermal fluids to the second largest Proterozoic orogenic gold deposit, after Ashanti
clastic rocks, in a manner suggested for many Carlin deposits. in West Africa. It was mined from 1876 to 2001 by under-
Additionally, Wilde et al. (2001) suggested that the thermo- ground operations to depths of 2.5 km. Total production was
chemical sulfate reduction of nearby anhydrite-bearing Devo- 1237 t Au from ore averaging 8.34 g t1. The ore-hosting se-
nian evaporites provided a large H2S reservoir. quence consists of dominantly intracontinental rift basin sed-
At Muruntau, gold is commonly associated with pyrite, arse- iments that were deposited between c.2560 and 1870 Ma
nopyrite, and scheelite; the sulfides comprise 1.5–2 vol% of the (Caddey et al., 1991; Dahl et al., 2006) and metamorphosed
more altered rock. This giant hydrothermal system is geochemi- to upper greenschist to lower amphibolite facies. The deposit is
cally complex; there appears to be a zoning of stibnite- and essentially restricted to structurally deformed areas in the
cinnabar-rich mineralization surrounding the main gold ore c.2.0 Ga Homestake Formation (Hark et al., 2008), which is a
body, and there is a very silver-rich ore body adjacent and to the 125 m-thick carbonate–silicate–sulfide facies iron formation.
southeast of Muruntau (e.g., Myutenbai). Gold ores average Ore bodies occur as tabular replacement zones of pyrrhotite
39 ppm Bi, 30 ppm Te, 34 ppm Se, and 30 ppm W (Koneev and arsenopyrite within the chemically reactive BIF, which
et al., 2005). Traces of base metal sulfides, Bi-bearing phases, contains 15–35% total iron, and vein-hosting shear zones.
tellurides, and sulfosalts are common. The main stage of alteration Dating of ore-related arsenopyrite by Re–Os methods indicates
is commonly described as albite–quartz–biotite–pyrrhotite– gold deposition in the BIF at 1736  8 Ma (Morelli et al.,
tourmaline, with later sericite, chlorite, and minor carbonate 2010), but it is apparent that some gold was remobilized in
and K-feldspar (Bierlein and Wilde, 2010). Muruntau is the only response to Cretaceous magmatism.
major orogenic gold deposit known to also produce platinum. Caddey et al. (1991) summarized geochemical patterns in
However, the recovered platinum is not associated with the the Homestake Formation. Sulfur, CO2, Ag, As, and Au were
gold-forming event. Rather, the platinum-group elements are enriched in the ores; this is consistent with the observed hy-
locally enriched in the carbonaceous host rocks at levels of drothermal siderite and sulfide formation in and surrounding
0.30 ppm Pd and 0.61 ppm Pt (Yermolayev et al., 1995). the ore bodies. Such sulfidic replacement of Fe-rich layers in
Ore-related fluid inclusions for Muruntau are of low to BIF characterizes not only Homestake, but BIF-hosted orogenic
moderate salinity, are CO2-rich with significant CH4 and N2, gold deposits in general (e.g., Geita, Morro Velho). Less com-
and show evidence of fluid immiscibility (Berger et al., 1994; monly, enrichments in Co, Cr, Cu, Ni, Pb, Se, Te, V, and Zn
Graupner et al., 1998, 2001; Zairi and Kurbanov, 1992). Most have been measured. Unlike most orogenic gold deposits,
homogenization temperatures suggest ore formation some- there is no increase in W concentrations in the ore zones.
where between 270 and 400  C, although it is unclear specifi- Gold:silver ratio is 5:1, and silver was also recovered during
cally from what part of this range are the most reliable mining. Trace element and REE data, in conjunction with Pb
estimates and good pressure estimates are lacking. At the isotope data for ore-stage sulfides, have shown the replacement
silver-rich Myutenbai ore body, fluid inclusion temperatures style ore and the shear zone-related ore likely has a similar
and compositions are similar to those from the Muruntau ores, origin (Frei et al., 2009). Sulfur isotope @ 34S data range from
but there is no evidence of immiscibility (Graupner et al., 5.6 to 9.8% for ore-related sulfides within the ore bodies (Rye
2001). Veins similar to those hosting ores, but distal to the and Rye, 1974), which were interpreted to indicate a seawater
Muruntau ore zones, are also CO2-rich, but lack elevated CH4, sulfate source for the sulfur. Rye and Rye (1974) also showed
N2, and H2S compositions (Wilde et al., 2001). that, locally, sulfur values of sulfides decreased toward the ore
Oxygen isotope values for ore-bearing quartz vein types bodies within the Homestake Formation. Data were inter-
range from 8 to 15%, with fluid values calculated at mainly preted as consistent with a syngenetic model for a gold source,
400  C between 4 and 9 per mil (Berger et al., 1994). Gold- although more recent structural (Caddey et al., 1991) and
related pyrite and arsenopyrite @ 34S values range between 2.8 geochronologic data (Morelli et al., 2010) confirm an epige-
and 5.5% (Berger et al., 1994). Helium in arsenopyrite indi- netic formation. Oxygen isotope @ 18O data for hydrothermal
cates a mainly crustal origin and a much lesser mantle source, quartz in the ore bodies range from 11.7 to 15.9%. Rye and
418 Geochemistry of Hydrothermal Gold Deposits

Rye (1974) suggested a trend in these data, with isotopically in dolerite surrounding the Oroya veinlets are altered
lighter measurements in areas of the deposit that were of to vanadium mica–ankerite–quartz–siderite–hematite–pyrite–
highest metamorphic grade. They also report dD values of 56 telluride (Bateman et al., 2001; White et al., 2003).
to 112% and d13C values of 1.3 to 8.4%, but these data Ho et al. (1990a, 1990b) interpreted the Fimiston ores to
are from fluid inclusions in crushed vein material, and the have formed from an aqueous-carbonic fluid with 15–25 mol%
meaning of such bulk extraction measurements is questionable. CO2 and a salinity of <5.5 wt% NaCl (equiv.). The Oroya style
Lead isotopes of sulfide minerals have been interpreted to of mineralization showed a significant methane component in
indicate an epigenetic mixture of magmatic and host-rock the hydrothermal fluids, which was interpreted to reflect the
leads (Frei et al., 2009), although the 1713  10 Ma age for interaction with lesser carbonaceous metasedimentary rocks
local magmatism postdates the Re–Os mineralization age (see within the overall metamafic sequence. Temperature-corrected
preceding text). fluid inclusion trapping pressures were estimated to be 1.5–4 kb
Fluid inclusion data were determined for arsenopyrite- at about 280–330  C. Clout (1989) suggested conflicting P–T
bearing quartz veins (Groen et al., 1989; Kath, 1990; Rye and estimates of ore formation of 100–260 bars and 170–250  C,
Rye, 1974). Two- and three-phase H2O–CO2 fluid inclusions also from fluid inclusion interpretations, but such shallow pres-
were related to the ore-forming event. A fluid inclusion ho- sures are atypical of most orogenic gold deposits (e.g., Groves
mogenization temperature mode of 248–274  C, with trap- et al., 1998), and the low temperatures are not consistent with
ping pressure estimates of 3.5–4 kb from metamorphic the mineral assemblages (e.g., White et al., 2003). Log fS2 was
conditions, yields trapping temperature estimates of 400– estimated as 12.6 to 5.5 and log fTe2 at 11.4 to 6.8 for the
450  C. Microthermometric data indicate as much as 8 mol% ore-forming fluids (Shackleton et al., 2003). Most vein quartz
CH4 in a fluid with salinities between 3.8 and 6.5 equiv. wt% measurements of @ 18O ranged from 14 to 15% with calculated
NaCl (equiv.). fluids of 6.6  2.0%, and dD values for hydrothermal muscovite
were generally 25 to 50% (Clout, 1989; Golding et al.,
13.15.4.3.3 Late Archean Greenstone-hosted Golden Mile 1990). The widespread hematitic alteration, coupled with the
The Golden Mile deposit in the Yilgarn craton of Western anomalously light d34S (2 to 10%) for Archean orogenic
Australia is the largest Archean orogenic gold deposit. It was gold pyrites (Phillips et al., 1986), has led to controversies in the
mined from the late 1890s through the late 1980s by under- ore genesis model in trying to explain the relatively oxidized
ground operations and with ores averaging about 11.5 g t1 hydrothermal fluids. Many workers have stressed a mixing of
Au. Since 1989, gold has been recovered from the giant fluids or two distinct fluids to help explain assemblages with
Fimiston open-pit and has averaged about 1.5–2 g t1 Au. coexisting hematite, magnetite, siderite, ankerite, pyrite, and
Total production has exceeded 1600 t Au and significant anhydrite (e.g., Neumayr et al., 2008). However, thermody-
resources remain. The deposit occurs along the regional namic modeling by Evans et al. (2006) shows how hematite
Boulder-Lefroy fault system. The greenstone belt volcanic- precipitation and a sulfate-bearing fluid could form via reaction
sedimentary rock succession is c.2.69 Ga in age and was in- of a typical reduced orogenic gold hydrothermal fluid with the
truded by the ore-hosting Golden Mile dolerite (diabase) dike Golden Mile dolerite.
at c.2675 Ma. The absolute age of gold formation has been
controversial, but recent data suggest that the ductile–brittle
Fimiston-type replacement/shear mineralization, the brittle– 13.15.5 Summary and Conclusions
ductile Oroya-type veinlet mineralization, and the brittle Mt.
Charlotte-type stockwork mineralization formed relatively Gold is transported by aqueous complexes formed with dis-
coeval at c.2640 Ma (Vielreicher et al., 2010). The reason for solved H2S in the ore-forming fluids in all of the ores discussed
the large size of the Golden Mile has been suggested to be the here. The H2S can come from igneous rocks, magmas, rocks
combination of favorable structural and geologic characteris- undergoing metamorphism, and even sedimentary rocks.
tics (Phillips et al., 1996), with the total gold and volatile Other elements with strong affinity to make aqueous com-
content of the ore bodies able to be explained by devolatiliza- plexes with H2S are transported along with gold, particularly
tion of a greenstone block 8 km  8 km  5 km in volume As, but also Hg, Sb, Se, Te, Tl, and locally Ag. Silver may
(Phillips et al., 1987). The gold ore is defined by an Ag–Au– decouple from these by forming earlier Ag–S minerals during
Te–W geochemical signature, with wall rock enrichments in transport or perhaps remains in the solution as stable Ag–Cl
CO2, K, S. B, V, Ba, As, Sb, and Rb (Bateman et al., 2001; complexes in some settings. Temperatures, salinities, of ore-
Phillips and Gibb, 1993). Pyrite is the dominant sulfide forming fluids in all ores discussed here are generally similar,
phase in the ores, but there are minor amounts of arsenopyrite, and all can locally be enriched in CO2. Orogenic gold deposits
tetrahedrite, and base metal sulfides. About 25% of the ores are formed at higher temperature and deeper than the other
hosted by a wide variety of telluride minerals (e.g., Shackleton deposit types have higher dissolved CO2 contents, and the
et al., 2003), with Te concentrations locally exceeding pH of the ore-forming fluids may be different from CTD and
1000 ppm. Gold fineness throughout the deposit is >900  C. epithermal deposits. Some combination of boiling, cooling,
Silver and Sb both increase relative to gold with a depth and fluid mixing cause deposition of most epithermal ores,
(Bateman et al., 2001), which is atypical of most hydrothermal whereas reactions with Fe-rich carbonate host rocks are most
systems. Within a few meters of the Fimiston lodes, high-grade important for CTD, whereas pressure drops (typically caused
altered dolerite is dominated by an ankerite–siderite–sericite– by hydrofracturing) are important in orogenic deposits. Per-
quartz–hematite–pyrite–telluride  albite, tourmaline, magne- haps gold nanoparticles can form in any of these deposits (less
tite, and anhydrite assemblage. Similar high-grade ore zones likely in orogenic?) when dissolved contents are elevated, and
 Geochemistry of Hydrothermal Gold Deposits 419

they are physically transported to the site of deposition. How- Castor SB, Boden DR, Henry CD, et al. (2003) The Tuscarora Au–Ag district: Eocene
ever, evidence of this is preserved only in the highest grade volcanic-hosted epithermal deposits in the Carlin gold region, Nevada. Economic
Geology 98: 339–366.
deposits, and the formation of nanoparticles implies a gold-
Chamberlain VE and Lambert RJ (1994) Lead isotopes and the sources of the Columbia
rich fluid evolves to form them along its flow path. River Basalt Group. Journal of Geophysical Research 99: 11805–11817.
Chang Z, Large RR, and Maslennikov V (2008) Sulfur isotopes in sediment-hosted
orogenic gold deposits: Evidence for an early timing and a seawater sulfur source.
Geology 36: 971–974.
Acknowledgments Chen M, Mao J, Qu W, et al. (2007) Re–Os dating of Arsenian pyrites from the Lannigou
gold deposit, Zhenfeng, Guizhou Province, and its geological significances.
This research was supported in part by funding from the US Dizhi Lunping Geological Review 53(3): 0371–5736.
Chesley JT and Ruiz J (1998) Crust–mantle interaction in large igneous provinces:
National Science Foundation to J. Saunders (EAR-0838208), Implications from the Re–Os isotope systematics in large igneous provinces:
US Geological Survey’s Mineral Resource External Research Implications from the Re–Os isotope systematics of the Columbia River flood
Program (MRERP #05HQGR0153). Thanks to Samuel B. basalts. Earth and Planetary Science Letters 154: 1–11.
Romberger and Erin E. Marsh who reviewed and edited an Chitwood RA (2012) Geochemistry and Mineralogy of the Eastern Ag-rich Epithermal
Veins in the Midas District, Nevada, USA, 117 p. Unpublished MS Thesis, Auburn
earlier version of this paper. The authors also thank Newmont
University.
and Barrick mining companies for allowing access to their data Chouinard A, Paquette J, and Williams-Jones A (2005a) Crystallographic controls on
and some of the ores discussed here. trace-element incorporation in auriferous pyrite from the Pascua epithermal
high-sulfidation deposit, Chile–Argentina. The Canadian Mineralogist
43: 951–963.
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