The following is an updated and expanded version of a relatively recent review paper on skarns (Meinert (1992)) and contains numerous links to WWW links, photographs, and examples of individual skarn deposits which are not in the published version.
This was translated thusly by Torbjorn Bergman of the University of Stockholm (1992, written communication), "As subordinate layers in the feldspar-poor felsic volcanic rocks, there appear peculiar dark rocks which also are the ore's host rock. These rocks are in the Persberg area denoted 'skarn', a word which likely can be used as a collective term for all such odd rocks occurring alongside the ores." Tornebohm goes on to describe garnet-rich "brunskarn" (brown skarn) and pyroxene-rich "gronskarn" (green skarn).
For other pre-1970 publications concerning skarn deposits, Burt (1982) provides a very useful annotated historical bibliography. Einaudi et al. (1981), a fairly exhaustive review paper on skarn deposits, is a good source of other references and definitions, some of which are summarized below. Another source of historical information is the recently published field trip guide to the classic Banat region of Romania. Additional references on skarn deposits are available here.
mineralogy includes a wide variety of calc-silicate and associated minerals but usually is dominated by garnet and pyroxene.
Skarns can be subdivided according to several criteria. Exoskarn and endoskarn are common terms used to indicate a sedimentary or igneous protolith, respectively. Magnesian and calcic skarn can be used to describe the dominant composition of the protolith and resulting skarn minerals. Such terms can be combined, as in the case of a magnesian exoskarn which contains forsterite-diopside skarn formed from dolostone.
Calc-silicate hornfels is a descriptive term often used for the relatively fine-grained calc-silicate rocks that result from metamorphism of impure carbonate units such as silty limestone or calcareous shale. Reaction skarns can form from isochemical metamorphism of thinly interlayered shale and carbonate units where metasomatic transfer of components between adjacent lithologies may occur on a small scale (perhaps centimetres) (e.g. Vidale, 1969; Zarayskiy et al., 1987). Skarnoid is a descriptive term for calc-silicate rocks which are relatively fine-grained, iron-poor, and which reflect, at least in part, the compositional control of the protolith (Korzkinskii, 1948; Zharikov, 1970). Genetically, skarnoid is intermediate between a purely metamorphic hornfels and a purely metasomatic, coarse-grained skarn.
For all of the preceding terms, the composition and texture of the protolith tend to control the composition and texture of the resulting skarn. In contrast, most economically important skarn deposits result from large scale metasomatic transfer, where fluid composition controls the resulting skarn and ore mineralogy. This is the mental image that most people share of a "classic" skarn deposit. Ironically, in the "classic" skarn locality described by Tornebohm at Persberg, skarn has developed during regional metamorphism of a mostly calcareous Proterozoic iron formation. This reinforces the importance of Einaudi et al.'s (1981) warning that the words "skarn" and "skarn deposits" be used strictly in a descriptive sense, based upon documented mineralogy, and free of genetic interpretations.
Not all skarns have economic mineralization; skarns which contain ore are called skarn deposits. In most large skarn deposits, skarn and ore minerals result from the same hydrothermal system even though there may be significant differences in the time/space distribution of these minerals on a local scale. Although rare, it is also possible to form skarn by metamorphism of pre-existing ore deposits as has been suggested for Aguilar, Argentina (Gemmell et al., 1992), Franklin Furnace, USA (Johnson et al., 1990), and Broken Hill, Australia (Hodgson, 1975).
Although many skarn minerals are typical rock-forming minerals, some are less abundant and most have compositional variations which can yield significant information about the environment of formation (e.g. pyroxene - Takano, 1998; scapolite - Pan, 1998). Table 1 lists many of the common skarn minerals and their end member compositions. Some minerals, such as quartz and calcite, are present in almost all skarns. Other minerals, such as humite, periclase, phlogopite, talc, serpentine, and brucite are typical of magnesian skarns but are absent from most other skarn types. Additionally, there are many tin, boron, beryllium, and fluorine-bearing minerals which have very restricted, but locally important, parageneses.
The advent of modern analytical techniques, particularly the electron microprobe, makes it relatively easy to determine accurate mineral compositions and consequently, to use precise mineralogical names. However, mineralogical names should be used correctly so as not to imply more than is known about the mineral composition. For example, the sequence pyroxene, clinopyroxene, calcic clinopyroxene, diopsidic pyroxene, and diopside, are increasingly more specific terms. Unfortunately, it is all too common in the geologic literature for specific end member terms, such as diopside, to be used when all that is known about the mineral in question is that it might be pyroxene.
Zharikov (1970) was perhaps the first to describe systematic variations in skarn mineralogy among the major skarn classes. He used phase equilibria, mineral compatibilities, and compositional variations in solid solution series to describe and predict characteristic mineral assemblages for different skarn types. His observations have been extended by Burt (1972) and Einaudi et al. (1981) to include a wide variety of deposit types and the mineralogical variations between types. The minerals which are most useful for both classification and exploration are those, such as garnet, pyroxene, and amphibole, which are present in all skarn types and which show marked compositional variability. For example, the manganiferous pyroxene, johannsenite, is found almost exclusively in zinc skarns. Its presence, without much further supporting information, is definitive of this skarn type.
When compositional information is available, it is possible to denote a mineral's composition in terms of mole percent of the end members. For example, a pyroxene which contains 70 mole percent hedenbergite, 28 mole percent diopside, and 2 mole percent johannsenite could be referred to as Hd70Di28Jo2. In many skarn systems, variation in iron content is the most important parameter and thus, many minerals are described simply by their iron end member, e.g. Hd10 or Ad90. Large amounts of compositional information can be summarized graphically. Triangular plots commonly are used to express variations in compositionally complex minerals such as garnet and pyroxene.
Amphiboles are more difficult to portray graphically because they have structural as well as compositional variations. The main differences between amphiboles in different skarn types are variations in the amount of Fe, Mg, Mn, Ca, Al, Na, and K. Amphiboles from Au, W, and Sn skarns are progressively more aluminous (actinolite-hastingsite-hornblende), amphiboles from Cu, Mo, and Fe skarns are progressively more iron-rich in the tremolite-actinolite series, and amphiboles from zinc skarns are both Mn-rich and Ca-deficient, ranging from actinolite to dannemorite. For a specific skarn deposit or group of skarns, compositional variations in less common mineral phases, such as idocrase, bustamite, or olivine, may provide insight into zonation patterns or regional petrogenesis (e.g. Giere, 1986; Agrell and Charnely, 1987; Silva and Siriwardena, 1988; Benkerrou and Fonteilles, 1989).
For example, circulating diverse fluids through a fracture in a relatively simple carbonate protolith can result in several different reactions. Thus, the steep thermal gradients common in most plutonic environments, result in complex metamorphic aureoles complete with small-scale metasomatic transfer as evidenced by reaction skarns and skarnoid.
More complex metasomatic fluids, with the possible addition of magmatic components such as Fe, Si, Cu, etc. , produce a continuum between purely metamorphic and purely metasomatic processes. This early metamorphism and continued metasomatism at relatively high temperature (Wallmach and Hatton, 1989, describe temperatures > 1200C) are followed by retrograde alteration as temperatures decline. A link between space and time is a common theme in ore deposits and requires careful interpretation of features which may appear to occur only in a particular place (e.g. Barton et al., 1991).
One of the more fundamental controls on skarn size, geometry, and style of alteration is the depth of formation. Quantitative geobarometric studies typically use mineral equilibria (Anovitz and Essene, 1990), fluid inclusions (Guy et al., 1989) or a combination of such methods (Hames et al., 1989) to estimate the depth of metamorphism. Qualitative methods include stratigraphic or other geologic reconstructions and interpretation of igneous textures. Simple observations of chilled margins, porphyry groundmass grain size, pluton morphology, and presence of brecciation and brittle fracture allow field distinctions between relatively shallow and deep environments.
The effect of depth on metamorphism is largely a function of the ambient wall rock temperature prior to, during, and post intrusion. Assuming an average geothermal gradient for an orogenic zone of about 35C per kilometre (Blackwell et al., 1990), the ambient wall rock temperature prior to intrusion at 2 km would be 70C, whereas at 12 km it would be 420C. Thus, with the added heat flux provided by local igneous activity, the volume of rock affected by temperatures in the 400-700C range would be considerably larger and longer lived surrounding a deeper skarn than a shallower one. In addition, higher ambient temperatures could affect the crystallization history of a pluton as well as minimize the amount of retrograde alteration of skarn minerals.
At a depth of 12 km with ambient temperatures around 400C, skarn may not cool below garnet and pyroxene stability without subsequent uplift or other tectonic changes. The greater extent and intensity of metamorphism at depth can affect the permeability of host rocks and reduce the amount of carbonate available for reaction with metasomatic fluids. An extreme case is described by Dick and Hodgson (1982) at Cantung, Canada, where the "Swiss cheese limestone" was almost entirely converted to a heterogeneous calc-silicate hornfels during metamorphism prior to skarn formation. The skarn formed from the few remaining patches of limestone has some of the highest known grades of tungsten skarn ore in the world (Mathiason and Clark, 1982).
The depth of skarn formation also will affect the mechanical properties of the host rocks. In a deep skarn environment, rocks will tend to deform in a ductile manner rather than fracture. Intrusive contacts with sedimentary rocks at depth tend to be sub-parallel to bedding; either the pluton intrudes along bedding planes or the sedimentary rocks fold or flow until they are aligned with the intrusive contact. Examples of skarns for which depth estimates exceed 5-10 km include Pine Creek, California (Brown et al., 1985) and Osgood Mountains, Nevada (Taylor, 1976). In deposits such as these, where intrusive contacts are sub-parallel to bedding planes, skarn is usually confined to a narrow, but vertically extensive, zone. At Pine Creek skarn is typically less than 10 m wide but locally exceeds one kilometre in length and vertical extent (Newberry, 1982).
Thus, skarn formed at greater depths can be seen as a narrow rind of small size relative to the associated pluton and its metamorphic aureole. In contrast, host rocks at shallow depths will tend to deform by fracturing and faulting rather than folding. In most of the 13 relatively shallow skarn deposits reviewed by Einaudi (1982a), intrusive contacts are sharply discordant to bedding and skarn cuts across bedding and massively replaces favorable beds, equalling or exceeding the (exposed) size of the associated pluton. The strong hydrofracturing associated with shallow level intrusions greatly increases the permeability of the host rocks, not only for igneous-related metasomatic fluids, but also for later, possibly cooler, meteoric fluids (Shelton, 1983). The influx of meteoric water and the consequent destruction of skarn minerals during retrograde alteration is one of the distinctive features of skarn formation in a shallow environment.
The shallowest (and youngest) known skarns are presently forming in active geothermal systems (McDowell and Elders, 1980; Cavarretta et al., 1982; Cavarretta and Puxeddu, 1990) and hot spring vents on the seafloor (Zierenberg and Shanks, 1983). These skarns represent the distal expression of magmatic activity and exposed igneous rocks (in drill core) are dominantly thin dikes and sills with chilled margins and a very fine grained to aphanitic groundmass.
The degree to which a particular alteration stage is developed in a specific skarn will depend on the local geologic environment of formation. For example, metamorphism will likely be more extensive and higher grade around a skarn formed at relatively great crustal depths than one formed under shallower conditions. Conversely, retrograde alteration during cooling, and possible interaction with meteroric water, will be more intense in a skarn formed at relatively shallow depths in the earth's crust compared with one formed at greater depths. In the deeper skarns carbonate rocks may deform in a ductile manner rather than through brittle fracture, with bedding parallel to the intrusive contact; in shallower systems the reverse may be true. These differences in structural style will in turn affect the size and morphology of skarn. Thus, host rock composition, depth of formation, and structural setting will all cause variations from the idealized "classic" skarn model.
Groupings of skarn deposits can be based on descriptive features such as protolith composition, rock type, and dominant economic metal(s) as well as genetic features such as mechanism of fluid movement, temperature of formation, and extent of magmatic involvement. The general trend of modern authors is to adopt a descriptive skarn classification based upon the dominant economic metals and then to modify individual categories based upon compositional, tectonic, or genetic variations. This is similar to the classification of porphyry deposits into porphyry copper, porphyry molybdenum, and porphyry tin types; deposits which share many alteration and geochemical features but are, nevertheless, easily distinguishable. Seven major skarn types (Au, Cu, Fe, Mo, Sn, W, and Zn-Pb) have received significant modern study and several others (including F, C, Ba, Pt, U, REE) are locally important. In addition, skarns can be mined for industrial minerals such as garnet and wollastonite.
Major skarn types:
Calcic iron skarns in oceanic island arcs are associated with iron-rich plutons intruded into limestone and volcanic wall rocks. In some deposits, the amount of endoskarn may exceed exoskarn. Skarn minerals consist dominantly of garnet and pyroxene with lesser epidote, ilvaite, and actinolite; all are iron-rich (Purtov et al., 1989). Alteration of igneous rocks is common with widespread albite, orthoclase, and scapolite veins and replacements, in addition to endoskarn.
In contrast, magnesian iron skarns are associated with diverse plutons in a variety of tectonic settings; the unifying feature is that they all form from dolomitic wall rocks. In magnesian skarns, the main skarn minerals, such as forsterite, diopside, periclase, talc, and serpentine, do not contain much iron; thus, the available iron in solution tends to form magnetite rather than andradite or hedenbergite (e.g. Hall et al., 1989).
Overprinting of calcic skarn upon magnesian skarn is reported from many Russian deposits (Sokolov and Grigorev, 1977; Aksyuk and Zharikov, 1988). In addition, many other skarn types contain pockets of massive magnetite which may be mined for iron on a local scale (e.g. Fierro area, New Mexico, Hernon and Jones, 1968). Most of these occurrences form from dolomitic strata or from zones that have experienced prior magnesian metasomatism (e.g. Imai and Yamazaki, 1967).
Prior to the dramatic rise in the price of gold in the early 1970s, most gold produced from skarn deposits came as a byproduct of the mining of other metals, particularly Cu. The one notable exception was the Nickel Plate mine in the Hedley district, British Columbia, which had been mined for high grade gold in skarn from the turn of the century (Billingsley & Hume 1941). This deposit has been intensively studied (Ray et al. 1986b, 1988, 1993, 1995, 1996; Ettlinger 1990; Ettlinger et al. 1992; Ray & Dawson 1987, 1988, 1994) and has served as a de facto exploration model for gold skarn deposits in combination with the relatively similar Fortitude deposit in Nevada (Wotruba et al. 1988; Myers & Meinert 1991; Theodore & Hammarstrom 1991; Myers 1994). Subsequent recognition of similar Au skarn deposits includes: Andorra, Spain (Romer & Soler 1995); Beal, Montana (Wilkie 1996); Buffalo Valley, Nevada (Seedorf et al. 1991); Crown Jewel, Washington (Hickey 1992); Elkhorn, Montana (Everson & Read 1992); Junction Reefs, Australia (Gray et al. 1995); Marn, Yukon (Brown & Nesbitt 1985); Redline, Nevada (Theodore & Hammarstrom 1991); Ximena, Ecuador (Paladines & Rosero 1996).
Numerous other gold skarn deposits have been discovered in the past several decades which differ in important ways from the Hedley-Fortitude model. For example, some are magnesian skarns (Butte Highlands, Montana, Ettlinger et al. 1996; Marvel Loch, Australia, Mueller 1991, Mueller et al. 1991), some are magnetite-dominant (Bermejal, Mexico, de la Garza et al. 1996; Key East, Washington, Lowe 1998), some are garnet-dominant and relatively oxidized (Ban Na Lom, Thailand, Pisutha-Arnond et al. 1984; McCoy, Nevada, Brooks 1994; Nambija, Ecuador, Hammarstrom 1992; Red Dome, Australia, Ewers & Sun 1989; Wabu, Irian Jaya, Allen et al. 1998), and some occur in iron-rich rocks in regional metamorphic terrains (Lucky Draw, Australia, Sheppard et al. 1995; Lupin, Northwest Territories, Lhotka & Nesbitt 1989; Mallapakonda and Oriental, India, Siddaiah & Rajamani 1989; Navachab, Namibia, Noertemann 1997, Moore and Jacob, 1998; Nevoria, Australia, Mueller 1997; Tillicum, British Columbia, Ray et al. 1986a; Peterson 1996). Reviews of gold-bearing skarns that contain useful background data include: Yakrushev (1972), Meinert (1989, 1998), Ray et al. (1990), Theodore et al. (1991), and Ray & Webster (1991, 1995).
The term "gold skarn" is used here in the economic sense suggested by Einaudi et al. (1981) and refers to ore deposits that are mined solely or predominantly for gold and which exhibit calc-silicate alteration, usually dominated by garnet and pyroxene, that is related to mineralization. This usage excludes deposits such as Big Gossan (Meinert et al., 1997) that contain substantial gold (>1 million ounces and > 1 g/t Au), but which are mined primarily for other commodities such as copper. It also excludes deposits such as the Veselyi Mine in the Siniukhinskoe District, CIS where gold was high-graded from a Cu-Au skarn system due to socioeconomic considerations, but which would have been mined for Cu-Au in most other societies (Ettlinger & Meinert 1992). Conversely, this definition includes deposits such as Bermejal (de la Garza et al. 1996) and Key East (Lowe & Larson 1996; Lowe 1998) that contain large amounts of other metals (such as Fe in the form of magnetite) that are not mined.
Reduced Gold Skarns
The highest grade (5-15 g/t Au) gold skarn deposits are relatively reduced, are mined solely for their gold content, lack economic concentrations of other metals, and have a distinctive Au-Bi-Te-As geochemical association. Most high-grade gold skarns are associated with reduced (ilmenite-bearing, Fe2O3/( Fe2O3+FeO) << 0.75) diorite-granodiorite plutons and dike/sill complexes. They typically occur in clastic-rich protoliths rather than pure limestone and skarn alteration of dikes, sills, and volcaniclastic units is common. Reduced gold skarns are dominated by iron-rich pyroxene (typically > Hd50), but proximal zones can contain abundant intermediate grandite garnet. Other common minerals include K-feldspar, scapolite, vesuvianite, apatite, and amphibole. Distal/early zones contain biotite+K-feldspar hornfels, that can extend for 100s of meters beyond massive skarn. Due to the clastic-rich, carbonaceous nature of the sedimentary rocks in these deposits, most skarn is relatively fine-grained.
Hedley District, British Columbia
The Nickel Plate mine in the Hedley district, British Columbia is the largest and highest grade gold skarn in Canada. Discontinuous production from 1904 until the mine closed in 1995 was 13.4 million tons averaging 5.3 g/t Au, 1.3 g/t Ag, and 0.02% Cu (Ray et al. 1996). Of this, more than 3 million tons of ore was mined underground at an even higher grade, averaging 14 g/t Au. Skarn formed in dominantly clastic rocks of the upper Triassic Nicola Group, that is part of the allochthonous Quesnel Terrane of the Intermontane Belt. Skarn is spatially and genetically associated with the dioritic Hedley intrusions, that comprise the Toronto Stock and a series of dikes and sills, many of which exhibit strong endoskarn alteration with abundant pyroxene, biotite, garnet, amphibole, and K-feldspar. Dating of these intrusions suggests an age range of 194-219 Ma (Ray & Dawson 1994). The Toronto Stock is a very reduced ilmenite-bearing intrusion with an average Fe2O3/( Fe2O3+FeO) value of 0.15, the lowest of any gold skarn (Ray et al. 1995) and the lowest of any major skarn class (Meinert 1995).
As first recognized by Billingsley & Hume (1941), skarn is zoned in both space and time relative to the Toronto Stock and associated dikes and sills. The earliest and most distal alteration is a fine-grained biotite hornfels that affects both clastic rocks and some of the early sills (Ray et al. 1988). With time and proximity to massive skarn, biotite occurs with K-feldspar and pyroxene and is slightly coarser grained (Ettlinger 1990). This forms an aureole around the massive garnet-pyroxene skarn that is zoned from garnet > pyroxene near the Toronto Stock to pyroxene-dominant (garnet:pyroxene <0.1) skarn in distal ore zones (Ettlinger et al. 1992). Garnet is intermediate grandite in composition whereas pyroxene is relatively iron-rich. The most iron-rich garnet (Ad73-82) occurs in distal ore zones (Ettlinger et al. 1992), whereas pyroxene composition changes systematically away from the Toronto Stock and the larger dikes, becoming more iron-rich and slightly more manganese-rich (Fig. 6). The sulfide minerals associated with garnet and pyroxene skarn are dominantly arsenopyrite, loellingite, and pyrrhotite. Other sulfides, in decreasing order of abundance, are chalcopyrite, pyrite, sphalerite, hedleyite (Bi2+XTe1-X), native bismuth, gold, galena, and maldonite (Au2Bi) (Ettlinger 1990). This latter group of minerals is mostly associated with lower temperature alteration including amphibole, ferroan wollastonite, scapolite, and prehnite. The scapolite and some amphiboles are unusually chlorine-rich and this feature has been suggested as an exploration guide to gold-rich systems (Pan 1998; Pan et al. 1994).
Garnet-pyroxene skarn at the Nickel Plate mine is estimated to have formed at a depth of 5 km and at an average temperature of 460-480°C, although fluid inclusions in some garnet and pyroxene samples homogenized at temperatures above 600°C (Ettlinger 1990). The presence of loellingite in addition to arsenopyrite also is suggestive of higher temperatures at a given sulfur fugacity (Kretschmar & Scott 1976, Heinrich & Eadington 1986). Salinities of garnet and pyroxene fluid inclusions average 18.3 and 9.7 wt. percent NaCl equivalent, respectively, with sparse halite daughter minerals in garnet, pyroxene, and quartz yielding a maximum of 37.9 wt. percent NaCl equivalent. Such high-salinity, high-temperature fluids would be capable of transporting significant gold as chloride complexes (Gammons & Williams-Jones 1995, 1997).
Fortitude Deposit, Battle Mountain District, Nevada
The Fortitude deposit is located in the Battle Mountain District of north-central Nevada and produced 77 Mt of Au from 10.9 Mt of ore at an average grade of 7.1 g/t Au (Doebrich & Theodore 1996). The Battle Mountain District contains several different skarn types ranging from a typical copper skarn with low gold grades, called the West Ore Body, to gold-rich, copper-poor skarns such as the Upper and Lower Fortitude ore bodies. The West ore body is adjacent to the 38-38.5 Ma (Theodore et al. 1978, K-Ar on primary biotite) Copper Canyon granodiorite porphyry. The skarns nearest the intrusive body are dominated by garnet with minor pyroxene and are typically high in copper with low concentrations of gold. The Fortitude deposit, like most high-grade Au skarns, has an unusual reduced skarn mineralogy and trace-element signature (Au-Bi-Te-As) that distinguishes it from most other skarn types. The reduced skarn mineralogy reflects the reduced nature of the associated Copper Canyon granodiorite [Fe2O3/( Fe2O3+FeO) < 0.5] that is quite distinct from typical oxidized porphyry copper deposit plutons. The more distal skarns contain more pyroxene than garnet and contain the highest concentrations of gold in the district (Myers 1994). An extensive biotite+diopsidic pyroxene+K-feldspar alteration halo surrounds the main skarn zone and extends up to 3 km from the Copper Canyon stock (Theodore & Blake 1975). Stylolites and other fluid escape structures are present in carbonate beyond the limit of calc-silicate formation.
Like much of central Nevada, detailed geologic relations in the Battle Mountain District are complicated by numerous thrust faults (De Witt, Golconda, and Roberts Mountain) that have juxtaposed a wide variety of rock types, all of which have been affected by hydrothermal alteration (Blake et al. 1984). Host rocks range from lower Cambrian through Permian and consist of sandstone, arkose, shale, chert, argillite, greenstone, limestone, and quartzite (Blake et al. 1984). Most skarn occurs in the basal mid-Pennsylvanian Battle Formation, the Pennsylvanian to Permian Antler Peak Limestone, and the Permian Edna Mountain Formation. The main mineralized unit is the Antler Peak Limestone that consists of well bedded dark gray limestone and minor chert nodules (Theodore & Blake 1978).
Skarn mineralogy has been investigated for most of the deposits in the Battle Mountain district (Theodore & Hammarstrom 1991; Doebrich et al. 1996). For example, the West Ore Body occurs mostly in the Antler Peak Limestone and is a typical copper skarn with a prograde skarn assemblage of grandite garnet (Ad39-99) + diopsidic pyroxene (Hd20-50) and a retrograde assemblage of actinolite + epidote + K-feldspar (Theodore & Blake 1978). Pyrrhotite, pyrite, chalcopyrite, and marcasite are the main sulfide minerals. The Fortitude deposit also occurs in the Antler Peak Sequence, primarily in the Antler Peak Limestone and the Battle Formation, and can be subdivided into an upper and lower ore body due to offset across the Virgin Fault. The Fortitude ore body contains a prograde skarn assemblage of hedenbergitic pyroxene (Hd20-95, Jo<10) + andraditic garnet, with only minor retrograde alteration to epidote + actinolite + chlorite. Pyroxene shows a general increase in iron content moving toward the marble front (Hd30 proximal to the Copper Canyon stock and Hd>90 at the distal skarn fringe). This trend is mimicked by a Mn enrichment moving towards the marble front, with the pyroxene nearest the intrusion rarely exceeding 3% Jo (except for crosscutting veinlets and crystal rims), whereas the pyroxene near the marble front is generally between 2-8% Jo (Myers 1994).
The distribution of most metals parallels the skarn zonation in the Antler Peak Limestone. Copper is highest in garnet-rich skarn near the intrusive contact, whereas gold is concentrated in pyroxene-dominant skarn, particularly where the pyroxene is iron-rich (>Hd50). Silver has elevated concentrations proximal to the stock and at the distal skarn front, beyond the main Fortitude gold zone (Kotlyar et al. 1998). The skarn system contains several sulfide species including chalcopyrite, pyrite, pyrrhotite, arsenopyrite, marcasite, sphalerite, and galena, that occur roughly in the order listed from intrusion to marble. Arsenopyrite is locally massive and native Bi is commonly visible in hand specimens. Native gold occurs at grain boundaries between skarn and sulfide minerals indicating a possible reaction relationship (Wotruba et al. 1988). In general, gold is associated with native bismuth, hedleyite, pearcite, and stannite. Trace elements are also zoned within the system with anomalous Co, Mo, Cr, and Ni in proximal zones and anomalous As, Bi, Cd, Mn, Pb, Zn, Sb, and Hg in distal zones.
Fluid inclusion work shows that the skarn formed at relatively high temperatures (300->550°C) that parallel fluid inclusion homogenization temperatures measured in the adjacent Virgin dike apophysis of the Copper Canyon granodiorite (Myers 1994). The distribution of measured fluid inclusion temperatures parallels the skarn zonation. Garnet closest to the main stock (drill hole #500) ranges from 360-590°C. More distal garnet and pyroxene (drill holes 2723 and 1997) range from 380-440°C and 320-430°C, respectively and the most distal (and iron-rich) pyroxene (drill hole 1994) ranges from 350-400°C (Myers 1994). In addition, high salinity conditions have been documented, with multiple daughter minerals in fluid inclusions identified by SEM and STEM analysis. Limited fluid inclusion measurements indicate pyroxene skarn had salinities of 25-44 wt. percent NaCl equivalent. Based upon limited evidence for boiling, Myers (1994) estimated a formation pressure of 0.4 kb (40 MPa) for the Fortitude system, in close agreement with the stratigraphic estimate of 1.5 km and a pressure of 375 bars (37.5 MPa) by Theodore & Blake (1975). As at Hedley, the presence of high-salinity, high-temperature fluids at Fortitude suggests gold transport by chloride complexes.
Similar zonation occurs in d18O and d13C values that indicate progressive reaction of a magmatic fluid with isotopically heavy carbonate wallrocks, as summarized by Zimmerman et al. (1992) and Myers (1994). Skarn garnets are progressively enriched in d18O outward from the Copper Canyon stock with garnet d18O values of 6.9 per mil in the proximal skarn and values as high as 8.2 per mil in distal skarn. Pyroxene (d18O = 8.6 to 10.3 per mil), amphibole (d18O = 8.6 to 9.2 per mil), and quartz (d18O = 11.4 to 13.2 per mil) are less systematic, but in each case the highest d18O values are most distal to the granodiorite stock. Skarn formation can be modelled as resulting from the progressive reaction of magmatic fluids with isotopically heavier carbonate wallrocks (d18O = 24.0 per mil). The variation in d13C values in calcite can also be explained by progressive reaction of magmatic fluids with carbonate wallrocks. d18O and d13C values decrease from unaltered limestone (d18O = 24.0 per mil, d13C = 2.4 per mil) to blocks of residual limestone in skarn (d18O = 15.4 to 19.3, d13C = -4.5 to 1.7 per mil) to calcite intergrown with skarn minerals (d18O = 11.8 to 13.1 per mil, d13C = -10.3 to -1.7 per mil). The absence of mineral phases with d18O less than magmatic values suggests that meteoric fluids (d18O <-10) probably did not play a significant role in the formation of this deposit (Zimmerman et al. 1992). This latter feature is consistent with the relatively small amount of retrograde alteration observed in the Fortitude skarn (Myers 1994).
Crown Jewel, Washington
Gold skarn mineralization at the Crown Jewel deposit was discovered in 1988 and current reserves stand at 8.7 Mt averaging 6.0 g/t Au, for a total of 52 t Au, although the deposit is still in a protracted permitting phase (Bob Derkey, Washington Department of Natural Resources, personal communication). Skarn is most closely associated with the Cretaceous (?) Buckhorn Mountain granodiorite and a series of granodiorite porphyry dikes that have been interpreted by Hickey (1990) as cogenetic. The main granodiorite pluton is quite reduced and contains primary ilmenite. It has a dioritic border phase that is more mafic and less silicic, but with similar alkalis relative to the central core. Hickey (1992) attributed the relatively high alkali content of the diorite border phase to alteration. The diorite is cut by garnet veins with pyroxene envelopes, although pervasive endoskarn alteration only occurs in some of the smaller dikes and sills.
The stratigraphy and structure of the host rocks at Crown Jewel are not well understood due to poor exposures and a regional metamorphic/shearing event that predates skarn formation (McMillen 1979). The shearing may be related to development of gneiss domes in the Okanogan highlands (Orr & Cheney 1987), although Hickey (1992) states that none of the skarn has been affected by shearing, e.g. there is no strain or deformation of skarn minerals. Rocks in the district that have been affected by alteration and mineralization can be divided into distinct groups including a lower unit containing calcareous siltstone, sandstone, and minor shale; a limestone that has been converted to marble; an upper unit containing shale, minor siltstone, and sandstone; and a distinctive chert pebble conglomerate (Hickey 1990). These units are thought to correlate with the Paleozoic Anarchist Formation. Structurally overlying the Anarchist is the Permo-Triassic Kobau Formation consisting of andesitic volcanic rocks with shale and volcaniclastic interbeds.
Distal alteration, especially in argillaceous and clastic units, consists of biotite and pyroxene hornfels. Closer to intrusive contacts or fluid pathways these minerals become coarser grained and pyroxene replaces the biotite. In more calcareous rocks and limestone, the early/distal biotite and pyroxene hornfels are replaced by garnet. Some of the rocks behaved in a brittle fashion following pyroxene formation such that veins and breccias are cemented by brown garnet. Close to intrusive contacts, limestone is completely replaced by massive garnet and magnetite. This zonation is mirrored by an iron enrichment in pyroxene, with the most distal pyroxene approaching pure hedenbergite in composition. Retrograde alteration at Crown Jewel is relatively coarse grained and consists of epidote, amphibole, zoisite, calcite, and quartz. Sulfides are associated with retrograde alteration and with massive magnetite. Magnetite-pyrrhotite occurs as veins cutting garnet close to the granodiorite, as well as massive replacement of marble. In places the magnetite is abundant enough to have been mined on a very small scale in the past, although it is not of economic importance at present.
Pyrrhotite is the most abundant sulfide mineral by far, reflecting the overall reduced nature of the protolith, pluton, and skarn mineralogy. Other sulfides include pyrite, marcasite, chalcopyrite, bismuthinite, cobaltite, native gold, native bismuth, and arsenopyrite (Hickey 1990). Arsenopyrite is only abundant in the relatively impermeable and brittle chert pebble conglomerate. As with most reduced gold skarns, bismuth minerals are strongly associated with gold mineralization. Crown Jewel may be unusual in that coarse grained bismuthinite is easily visible in drill core and is an excellent indicator of ore-grade gold (which is not visible at the hand specimen level). This bismuth-gold association is substantiated by assays of drill core composites.
Fluids associated with skarn formation and mineralization at Crown Jewel were high-temperature saline brines. Hickey (1990) reports abundant, large halite daughter minerals in plutonic quartz, but did not find daughter minerals in the very small inclusions present in skarn minerals. Primary fluid inclusions in pyroxene homogenized from 365-450°C, whereas those in garnet, homogenized from 300-370°C. Two salinity determinations from fluid inclusions in garnet yielded values of 19 and 22 eq. wt. % NaCl. Fluid inclusions in epidote and amphibole yielded slightly lower homogenization temperatures of 255-320°C and 315-350°C, respectively, for retrograde alteration. Based upon an assumed depth of 4 km at the time of intrusion and skarn formation, Hickey (1990) determined an average lithostatic pressure corrected temperature for garnet-pyroxene skarn of 465°C. Quartz veins that cut garnet-pyroxene skarn have similar homogenization temperatures with a wider range of salinity from 2-24 eq. wt. % NaCl.
The Elkhorn district in Montana contains a variety of reduced gold skarns related to mafic diorite stocks marginal to the Boulder Batholith. Individual skarn deposits include Carmody, Diamond Hill, Dolcoath, East Butte, Elkhorn, Heagan, Glory Hole, and Sourdough. Historic production from skarns in the Elkhorn district was 2.1 tons of Au as a byproduct of base-metal mining (Klepper et al. 1971). Recent exploration by several companies in the district has defined a combined resource of about 9 Mt averaging 4.8 g/t Au, based upon drilling of numerous discrete skarn zones (Everson & Read 1992 and unpublished abstracts). This represents a combined resource of more than 45 tons of gold contained in skarn.
The main phase of the Boulder Batholith is quartz monzonite dated at 75.7 ± 2.8 Ma (Everson & Read 1992). Satellite stocks at East Butte, Black Butte, and Cemetery Ridge stocks are dark, fine to medium-grained diorites which are similar in age to slightly older than the Boulder Batholith (Everson & Read 1992). These plutons have intruded a lower Paleozoic sequence including the Wolsey, Meagher, Park, Pilgrim, Maywood, Red Lion, Jefferson, Three Forks, and Madison formations. Near plutons, argillaceous rocks of the Park, Wolsey, and Three Forks formations have been converted to biotite, pyroxene, and calc-silicate hornfels, similar to that described at many other gold skarns, whereas the generally dolomitic carbonate units of the Meagher, Pilgrim, Maywood, Red Lion, Jefferson, and Madison formations have been recrystallized and locally silicified.
Skarn associated with the East Butte Diorite occurs as endoskarn in the diorite and as exoskarn in two stratabound units near the Wolsey-Meagher contact, which strikes NNW and dips 60-70°E (Everson & Read 1992). Exoskarn consists of dark green pyroxene and minor garnet. Pyrite, pyrrhotite, magnetite, and arsenopyrite occur disseminated in skarn, averaging 3-5%, and as massive replacement zones near the marble front. Minor phases recognized petrographically include marcasite, maldonite (Au2Bi), hedleyite (Bi14Te6), hessite (Ag2Te), gersdorfitte (NiAsS), and native bismuth (Meinert unpublished data). Retrograde alteration consists of amphibole, phlogopite, vesuvianite, and epidote. About half of the mineralization occurs as endoskarn alteration of the East Butte Diorite (Everson & Read 1992). Endoskarn consists of pyroxene, calcic plagioclase (close to pure anorthite), amphibole, titanite, and local veins of quartz-orthoclase.
Near the historic Carmody mine (Klepper et al. 1971), skarn associated with the East Butte Diorite occurs as a single stratabound layer in the Wolsey Formation. Skarn is presumed to have replaced a carbonate layer and is surrounded by biotite hornfels in the originally more argillaceous lithologies. Carmody mineralization is similar to the previously described East Butte mineralization except pyrrhotite is more abundant than pyrite, and both chalcopyrite and sphalerite occur in minor amounts.
Significantly different skarn mineralization occurs in the Sourdough zone northwest of East Butte near the historic Golden Curry mine (Roby et al. 1960). Sourdough skarn is spatially associated with both monzonite and diorite and occurs as endoskarn within the monzonite and as replacement of dolomitic rocks thought to be either Pilgrim or Jefferson Formation (Everson & Read 1992). Both endoskarn and exoskarn are pyroxene dominant with little garnet. Massive magnetite occurs at the marble front and within exoskarn associated with pyroxene, olivine, ludwigite-vonsenite ((Mg,Fe)BO5), and phlogopite. Retrograde alteration consists of abundant serpentine and tremolite.
Junction Reefs, Australia
Although large scale production is relatively recent, the gold skarns at Junction Reefs, New South Wales, Australia have been mined since 1876 with historical production of 1.1 tons of Au between 1876 and 1938 (Gray et al. 1995). Open pit mining began in 1988 and continues to the present, with total skarn reserves and production of 2.4 Mt with an average grade of 3.3 g/t Au, representing 7.7 tons of Au.
The protolith for skarn mineralization at Junction Reefs is a 39 m thick sequence of marine limestone, siltstone, and chert that occurs within the voluminous (>2500 m) Early Ordovician Coombing Formation consisting of massive volcanic graywacke, cherty shale, siltstone, and tuffaceous arenite (Gray et al. 1995). Like most turbidite sequences in an island arc setting, there are intercalations of volcanic flows and tuffs, but some workers regard the overall tectonic setting as one of shallow basins overlying a thin continental crust (Wyborn 1988). Intrusive into the Coombing Formation are a series of shoshonitic diorites, monzodiorites, monzonites, and quartz monzonites. In the Junction Reefs district numerous, locally interconnected, monzodiorite stocks, dikes, and sills were intruded between 430 and 440 Ma (Gray et al. 1995).
The Junction Reefs monzodiorite is surrounded by a zoned skarn system which has ore grade gold mineralization in the outer zones. Because most of the Coombing Formation consists of relatively unreactive siliciclastic rocks, skarn formation and mineralization are restricted to stratigraphic/structural windows of more calcareous rocks within the metamorphic/hydrothermal aureole of the monzodiorite and associated dikes and sills. However, as in many other gold skarn districts, the siliciclastic rocks have been converted to purple-brown biotite hornfels within 200 meters of the Junction Reefs monzodiorite (Gray et al. 1995). Closer to the intrusion and along bedding planes, fluids forming amphibole and pyroxene have infiltrated the rock, forming a green biotite-amphibole-pyroxene hornfels. This rock is not visually striking except when split open along bedding planes to expose radiating clusters of dark green amphibole and pyroxene crystals with interstitial diamond-shaped arsenopyrite up to 1 cm in length.
The skarn system is zoned around the Junction Reefs monzodiorite and three separate mines (Sheahan-Grants, Frenchmans, and Cornishmens) occur where calcareous rocks are exposed in the outer skarn zones. The inner most skarn zone (termed Zone 1) consists dominantly of pale green garnet, lesser quartz, and < 20% pyroxene. Both garnet and pyroxene range up to the pure Fe end members. Minor pyrite (< 5%) is disseminated in the garnet skarn and gold grades are low, averaging 0.1-0.2 g/t Au (Gray et al. 1995). In Zone 2 pyroxene is much more abundant than garnet and is slightly more iron-rich, on average, than in Zone 1, ranging from ferrosalite to hedenbergite (Hope 1990). Minor chlorite is reported with pyrrhotite and pyrite (Grant 1988). In Zone 3 the prograde pyroxene>>garnet skarn has been strongly overprinted by amphibole approaching ferrohastingsite in composition. Pyrrhotite is the dominant sulfide and is associated with amphibole, and to a lesser extent, with chlorite, calcite, and quartz. Gold reaches ore grade (>1.0 g/t Au) locally within Zone 3. Zone 4 hosts the vast bulk of ore grade mineralization. Remnant textures of garnet and pyroxene are present and rare small grains survive armored in quartz or sulfide, but most Zone 4 rocks are a dark green felted mass of chlorite, calcite, quartz, and sulfides. The dominant sulfide is pyrrhotite with lesser arsenopyrite, chalcopyrite, pyrite, and marcasite. Minor phases include native bismuth, maldonite, and an unidentified Au-Bi sulfide mineral. Zone 4 averages 10-20% sulfide and >80% massive sulfide occurs locally at the marble front. Gold is associated with sulfides and high concentrations of arsenopyrite (core assays range from 0.01-9.55% As) correlate with very high gold grades (Gray et al. 1995). Locally, there is a zone of wollastonite, vesuvianite, quartz ± grossularitic garnet at the marble front. This has been designated Zone 5, but is not as continuous as the other four zones.
Geochemically, skarn at Junction Reefs is anomalous in Au, As, Bi, Co, Fe, Pb, and Zn. As with many other gold skarns, the strongest correlation (r = 0.83) is between Au and Bi. Au and As are only moderately correlated (r = 0.58) and most other elements do not exhibit a systematic correlation with Au (Gray et al. 1995). Even in ore zones, Ag is very low, < 3 ppm. Fluid inclusions have not been examined in calc-silicate minerals from Junction Reefs. However, fluid inclusions in quartz and calcite have homogenization temperatures up to 345°C and 325°C, respectively and salt daughter minerals were observed in some inclusions, indicating at least some fluid salinities > 26 eq. wt. % NaCl (Grant 1988). These temperatures are in broad agreement with those determined for retrograde alteration at Junction Reefs from chlorite geothermometry (Grant 1988), thus indicating a minimum temperature for the system.
The Beal deposit is located approximately 26 km west-southwest of Butte, Montana and has proven and probable ore reserves of 14.8 Mt at an ore grade of 1.49 g/t Au, totaling 23.1 tons of Au. The deposit is hosted by late Cretaceous clastic, fluvial-deltaic sedimentary rocks of the Vaughn member of the Blackleaf Formation (Wilkie 1996). In the vicinity of the Beal mine, the Blackleaf Formation has been metamorphosed and metasomatized to a peak grade of pyroxene hornfels by diorite and granodiorite intrusions (74.8 + 2.8 m.y., K-Ar date on biotite) related to the Boulder Batholith (Hastings & Harrold 1988). A K-Ar date (71.7 + 2.6 m.y.) on adularia in a gold-bearing vein at Beal suggests that mineralization and intrusion are closely related (Hastings & Harrold 1988).
Structurally, the Beal deposit lies approximately three kilometers east of the western margin of the frontal fold and thrust belt of southwestern Montana. This zone is marked by a series of north-south trending thrust faults (Johnson, Spring, and Long Tom Thrusts), which juxtapose older (Paleozoic and Precambrian) rocks over the Cretaceous Blackleaf Formation (Ruppel et al. 1981). Thrusting predates the Beal deposit and is unrelated to mineralization. Numerous steep faults cut the Beal deposit, the most important being the German Gulch fault, Beal shear zone, and Gully fault. The Beal shear zone trends N80-85°W and dips 85-90°S, is locally mineralized, and was an important structural control for channeling hydrothermal fluids (Wilkie 1996).
All known mineralization at the Beal deposit occurs within the hornfels aureole of the granodiorite and diorite intrusions. Granodiorite of the Boulder Batholith crops out along the eastern edge of the mine area and numerous small dioritic stocks and dikes crop out near the margin of the batholith and within the open pit mine. Diorite samples from the pit are dark greenish-gray to greenish-black and consist of fine- to medium-grained plagioclase, biotite, amphibole, pyroxene, and K-spar. Opaque mineralogy consists of pyrite, chalcopyrite and magnetite-ilmenite intergrowths. The presence of ilmenite is indicative of a reduced magma chemistry and may be related to the gold content of this igneous system. All of the diorite exposures in the pit are intensely altered and contain up to 15-20% hydrothermal biotite/chlorite alteration of primary pyroxene, hornblende, biotite, and feldspar (Wilkie 1996).
Samples from traverses extending 3 km E-W perpendicular to the intrusive contact demonstrate mineralogical and temperature zonation outward from the pluton as follows: granodiorite --> pyroxene --> amphibole --> biotite --> white mica (Wilkie 1996). The width of the mineralogical zones is approximately constant throughout the area and within each zone mineral abundance decreases regularly (for a given protolith) with distance from the pluton. A marked exception is the high abundance of pyroxene near the Beal mine. Scapolite (containing 2-3 wt.% Cl) is also abundant in this area. Biotite, chlorite, and sericite geothermometers [models of McDowell and Elders (1980) and Walshe (1986)] indicate a temperature decrease away from the intrusive contact with a thermal anomaly coincident with the pyroxene-scapolite zone in the traverse which passes through the ore deposit (Wilkie 1996).
Pyroxene in pale green pyroxene hornfels occurs as <0.1 mm equant grains or clusters in the matrix among quartz grains. Most pyroxene compositions are from Hd18-40. Close to the Beal mine, pyroxene is slightly coarser grained, darker green, and more Fe-rich (Hd42-62). Coarser-grained pyroxene occurs along fractures and at fracture intersections as orbicular patches. Even in these occurrences pyroxene never constitutes a majority of the rock. Thus, alteration at Beal has far less pyroxene and overall is far less intense (garnet is absent) than any of the other reduced skarns considered in this review. This is considered to be a direct result of the carbonate-poor protolith and the relatively low temperature of skarn formation (Wilkie 1996).
Amphibole at Beal is compositionally more complex than pyroxene and ranges from calcic tremolite-actinolite to hornblende and ferrohastingsite. The iron-rich nature of the amphibole is surprising in light of the iron-poor nature of the host rocks. Locally, amphibole occurs as a replacement of pyroxene, but in most cases there is not textural or compositional evidence for pre-existing pyroxene (Wilkie 1996). Thus, amphibole at Beal appears to be a primary alteration phase. Strongly mineralized rocks can contain more than 50% amphibole and the amphibole typically is associated with K-feldspar, chlorite, and sulfides.
The main sulfide phases include pyrrhotite, pyrite, chalcopyrite, arsenopyrite, and minor tetrahedrite. Overall, Beal is a sulfide-poor system. Sulfide contents typically are less than 2% and many areas of ore grade material contain less than 1% total sulfide (Wilkie 1996). Local areas of > 5% arsenopyrite and/or pyrrhotite occur but are not always associated with high gold grades. Gold mineralization occurs as 1-5 micron grains disseminated in coarser-grained meta-sedimentary rocks and as flakes in quartz and quartz-sulfide-adularia veins (Hastings & Harrold 1988). The quartz-sulfide-adularia veins occur in steep structures within the pit and cut both the hornfels and plutons. Geochemical analyses of Cu, Bi, Zn, As, Au, and Ag in run-of-mine ore show that Au strongly correlates with Bi, weakly correlates with As and Cu, but does not correlate with other elements (Wilkie 1996).
Due to the very small grain size of most calc-silicate minerals at Beal, fluid inclusion studies have been limited. However, Wilkie (1996) measured fluid inclusions in quartz from several alteration styles at Beal. Measured homogenization temperatures range from 115-550°C with averages for L-V inclusions, L-V-salt, and vapor-rich inclusions of 344°, 312°, and 359°C, respectively. Salinities ranged from 4.7-42 eq. wt. % NaCl; the salinity of fluid inclusions with daughter minerals ranged from 34-42 eq. wt. % NaCl. Evidence for boiling in association with Au-Bi-Te minerals was documented at 340°C and salinities of 5-20 eq. wt. % NaCl. Similar temperatures were determined from mineral composition geothermometers, including arsenopyrite (408-428°C), biotite (362-400°C), and chlorite (314-378°C). In addition, both biotite and chlorite geothermometers indicate an overall temperature decrease away from the main intrusive contact.
d18O values for both coarser-grained (13-15.9) and finer-grained (12.9-15.5) siltstone samples at Beal increase away from the Boulder Batholith; coarser-grained samples have d18O values that are generally 0.5 higher than finer-grained samples for a given location Wilkie (1996). This distribution is a function of two interacting processes: 1) decreasing temperature and 2) decreasing water-rock ratio away from the contact. d 34S isotopic values are tightly clustered for a given mineral: pyrrhotite (7.7-9.3), pyrite (7.0-8.2), chalcopyrite (7.6-9.1), arsenopyrite (8.8-14.0), galena (9.2), and sphalerite (13.0). In most cases, fine-grained disseminated sulfides have lower d 34S isotopic values than vein sulfides and coarse-grained crystals. There is no evidence for a sedimentary sulfur reservoir; most samples are consistent with a dominantly magmatic source of sulfur (Wilkie & Meinert 1994).
Oxidized Gold Skarns
Whereas the "classic" gold skarn deposit is characterized by low garnet:pyroxene ratios, hedenbergitic pyroxene, and abundant sulfides dominated by pyrrhotite and arsenopyrite, several skarns have been mined for gold that have a very different mineralogy and mineralization style. These deposits have been classified by Brooks et al. (1991) as oxidized gold skarns. Their essential features include high garnet:pyroxene ratios, relatively Fe-poor garnet and pyroxene, low total sulfides, pyrite>pyyrhotite, and minor but ubiquitous occurrences of chalcopyrite, sphalerite, and galena. In addition, the highest gold grades are not associated with prograde garnet-pyroxene, but rather with later retrograde alteration including abundant K-feldspar (adularia) and quartz. Some of these deposits can be considered transitional to other types of gold mineralization such as epithermal deposits in which phase separation (boiling) can be an important precipitation mechanism (e.g. Hedenquist et al. 1996).
The McCoy gold skarn is only 45 km southwest of the reduced Fortitude gold skarn in northcentral Nevada, but differs dramatically in regards to the style of mineralization and wallrock alteration. The McCoy deposit contains 15.6 Mt of ore averaging 1.44 g/t Au and an additional 30,430 tonnes averaging 14.6 g/t Au that was mined underground (Brooks 1994). Production is from garnet-rich skarn surrounding the 39 Ma Brown stock, a reduced ilmenite-series, hypabyssal, hornblende-biotite granodiorite. Brooks (1994) subdivided the Brown stock into five petrologically distinct phases and invoked mixing of discrete magmas to yield individual intrusive phases. Importantly, there are systematic correlations between individual intrusive phases and the mineralogy and gold grade of associated skarn. The Brown stock is estimated to have intruded to within 1.3 km of the surface and this shallow emplacement is reflected by the multitude of dikes and sills found on the margins of the main stock. In addition, most of the early dikes and sills have been affected by garnet-pyroxene endoskarn.
Skarn at McCoy is zoned in both space and time. The earliest and most distal alteration is biotite and pyroxene hornfels. This results in a pale, fine-grained rock with original sedimentary layering still preserved. Overprinting this hornfels are veins and massive zones of garnet-dominant skarn. Typical garnet:pyroxene ratios are 3:1 to 20:1. Close to intrusive contacts, all the hornfels has been replaced and no trace of sedimentary bedding is left. Skarns closest to the main intrusion, called the West Contact and Peacock skarns, are the only skarns with significant pyroxene (>10%), and also the only pyroxene that is relatively coarse-grained and Fe-rich (up to Hd75). All other skarn at McCoy is garnet dominant and where pyroxene is present, it is diopsidic. Early garnet is Fe-poor and occurs as bedding replacements of argillaceous layers (skarnoid) and as cores to later metasomatic garnets, that are more Fe-rich. These compositional differences are important in that subsequent retrograde alteration selectively replaces certain stages and compositions of garnet and pyroxene (Brooks 1994). Sulfide minerals associated with prograde skarn include pyrrhotite, pyrite, sphalerite, galena, arsenopyrite, chalcopyrite, bornite, gold, hedleyite, native bismuth, and hessite (Brooks 1994).
Late garnet-pyroxene skarn coexists with or is overprinted by retrograde alteration consisting mainly of epidote-quartz-pyrite-K-feldspar. As previously described, grandite garnet is more susceptible to retrograde alteration than is andradite garnet. Biotite and chlorite occur instead of epidote in distal zones of retrograde alteration and where pyroxene was relatively abundant. Most economic gold mineralization is associated with retrograde alteration, particularly with quartz-pyrite-K-feldspar. The K-feldspar varies in color from pink to a pale tan and is similar to the adularia described from many epithermal deposits. The most intense quartz-pyrite-K-feldspar is spatially associated with a particular generation of dikes and sills called the Productive Series (Brooks 1994). However, quartz-pyrite-K-feldspar also replaces distal skarn and locally occurs as silicified pods in limestone beyond the limit of garnet-pyroxene alteration. This latter occurrence is similar to the jasperoids associated with some epithermal gold deposits, particularly Carlin-type deposits.
The fluids associated with prograde garnet and pyroxene at McCoy are high-temperature brines. Brooks (1994) reported homogenization temperatures in garnet ranging from 330-590°C with an average of 493±46°C. Measured salinities ranged up to 39.8 eq. wt. % NaCl. Homogenization temperatures in pyroxene range from 300-420°C and the average for proximal pyroxene is 398±14°C, whereas the average for distal pyroxene is 322±14°C. This spatial decrease in temperature is mirrored by a decrease in salinity. The salinity of fluid inclusions in proximal pyroxene ranges up to 35.3 eq. wt. % NaCl, whereas the maximum measured salinity in distal pyroxene is 22 eq. wt. % NaCl.
The fluids associated with retrograde alteration are slightly lower in temperature and salinity than those measured in prograde skarn, but are well above values typically reported for epithermal systems. Fluid inclusions in epidote (which replaces garnet) range from 360-450°C with salinities up to 28 eq. wt. % NaCl. Both the temperature and salinity of fluid inclusions in epidote are less than the values measured in garnet. Fluid inclusions were also measured in quartz and K-feldspar associated with retrograde alteration. Fluid inclusions in vein quartz range from 280-360°C with salinities from 11-19 eq. wt. % NaCl. Fluid inclusions in K-feldspar range from 160-390°C with salinities from 17-32 eq. wt. % NaCl.
Brooks (1994) estimated a pressure of 350 bars for skarn formation at McCoy and used this to determine an average pressure correction of 30°C for the measured homogenization temperatures. Collectively, these data indicate that prograde skarn formed at 330-620°C from brines with salinities up to 35 eq. wt. % NaCl. As temperatures declined, the early-formed garnet and pyroxene were altered to lower temperature assemblages including epidote, quartz, and K-feldspar. These minerals also formed from saline brines, but at temperatures 100-200°C lower than the prograde garnet and pyroxene.
Ecuador has two significant gold-bearing skarns, Ximena and Nambija. Ximena in west-central Ecuador is a typical reduced gold skarn similar to Hedley and Fortitude in North America. It has produced about 75,000 ounces of gold from alluvial fields developed from a pyroxene-dominant skarn. In contrast, Nambija in southeastern Ecuador is an oxidized gold skarn with similarities to the McCoy skarn in Nevada and Red Dome in Australia. Its mineralogy is dominated by grandite garnet and most production has come from alluvial workings and high-grading by local campesinos. Nambija may be best known for spectacular color photographs in the popular press that illustrate an ant swarm of human workers in the open pits reminiscent of gold rush days of previous centuries. Geologically, less is known about Nambija than most other gold skarn deposits due to the lack of organized mapping and the "unsettled" property ownership situation relative to the surface workings.
Nambija is one of a series of gold deposits in the southern portion of the Cordillera Real, a north-northeasterly trending belt of Cenozoic, Mesozoic, and Paleozoic rocks. The central part of this belt consists of Tertiary to recent volcanic rocks, with several active volcanoes. West of this volcanic belt is an accreted Cretaceous sequence of island arc and oceanic sedimentary, volcaniclastic, and volcanic rocks, that have been intruded by numerous Tertiary I-type, relatively mafic plutons. This belt hosts the Ximena gold skarn deposit. East of the central volcanic belt lies a series of Paleozoic metamorphic rocks and Mesozoic sedimentary-volcanic rocks, which have been deformed by a fold and thrust belt. Along the general contact between the Paleozoic and Mesozoic belts are several large Jurassic plutons and the Nambija deposit is located in a pendent in one of these batholiths.
On a regional scale, the Nambija district is dissected by west-verging, N10°E to N20°E thrust faults spaced approximately 10-30 km apart. The Nambija skarn deposits occur within metamorphosed Piuntza volcano-sedimentary rocks that occur as roof pendants in the 170 Ma Zamora batholith (Litherland et al. 1994). The Piuntza Unit is approximately 500 m thick and consists of sandstone, siltstone, limestone, tuff, and andesitic flows (Paladines & Rosero 1996). The Zamora batholith is an equigranular tonalite to granodiorite (Salazar 1988). Other igneous rocks that have been reported in the Nambija district include monzodiorite, monzonite, rhyodacite, syenite, and quartz-feldspar porphyry dikes and small stocks (Hammarstrom 1992; Paladines & Rosero 1996). However, most of these intrusions have been altered to K-feldspar, sericite, chlorite, and clay. Thus, the original compositions and ages of these intrusions are not well known.
Within the Nambija district, there are a series of gold-bearing skarns, that have been worked by local campesinos, including from north to south, Fortuna, Campana, Campanilla, Nambija, Guaysimi, and Sultana del Cóndor. Artisanal workings at Nambija are estimated to have produced 2 million ounces of gold and the current resource is estimated at 23 Mt (Mining Magazine 1990). Reported grades range from 14 to 84 g/t Au, with an average of 15-30 g/t Au (McKelvey 1991; Hammarstrom 1992). Campanilla and Campana are smaller but of similar grade (Mining Magazine 1990). Given the coarse grain size of the gold and the rudimentary nature of the alluvial and artisanal workings, all of the above tonnage and grade figures should be viewed with caution. Most skarn pockets and mineralized zones occur in a north-northeast structural corridor of breccias, veins, and shears that parallel the larger faults. This mineralized zone is 1.5 km long, 125 m wide, and dips 34°E within the pendent (Aguirre et al. 1985; McKelvey 1991). The highest-grade mineralization occurs at the intersection of these northerly structures and northeast striking faults. Where these intersecting fault zones cut skarn, the rock is dissected by parallel quartz stringers with native gold and few if any sulfide minerals (Aguirre et al. 1985). The fact that most of the mineralization and some of the skarn is structurally controlled and spatially associated with porphyritic rocks suggests that skarn formation and mineralization are not related to the main phase of the Zamora granodiorite. Instead, skarn formation appears to be associated with some of the younger porphyritic intrusions and mineralization is associated with quartz stringers that have a strong structural control.
There is a stock of quartz monzonite or rhyodacite porphyry at Nambija in the Tierrero 2 mine. The stock is surrounded by green garnet skarn with a zone of pink K-feldspar flooding and brecciation to the Southwest. The skarn is not sulfide-rich, but most samples contain minor pyrite, chalcopyrite, sphalerite and/or galena. In hand specimen, both the garnet and pyroxene are pale green in color. In addition, some of the garnet has pale brown and yellow hues as well. Such pale green-yellow garnet is typical of distal skarn and is similar to the garnet in many Zn skarns. In thin section, the garnet is strongly zoned as is typical of hydrothermal skarn garnet. There are discrete cores and rims to most grains indicating multiple pulses of hydrothermal fluid and in general, rims are more andraditic than cores, e.g. normal zoning. Almost all the garnet analyses reported from these rocks range from Ad21-72 except for a few distal samples with pure andradite. Although not highly anomalous, most of the garnets contain 0.5-1.5% MnO. This is slightly more spessartine component that would typically occur in Au skarn garnets. Otherwise, these intermediate grandite compositions are typical of Au skarns and would be quite unusual for most base metal skarn systems, including Fe, Cu, and Zn-Pb (Meinert 1992). In contrast, all the pyroxene is diopsidic and such iron-poor pyroxenes are atypical of Au skarns. The pyroxenes also are relatively manganese-rich (Hd16-34 Jo5-13), more than any other reported Au skarn, but significantly less than typical Zn skarns. The combination of high garnet:pyroxene ratios and both iron-poor garnet and pyroxene suggests that the Nambija system is both oxidized and iron-poor. This is consistent with the mineral abundances, compositions, and relative lack of iron sulfide minerals.
Gold at Nambija occurs with quartz veins spatially associated with garnet skarn. Some of the quartz veins have garnet envelopes indicating general contemporaneity with skarn formation. The fluid inclusions in the quartz are simple two phase inclusions. There are no daughter minerals, so the total salinity is < 26 wt. % NaCl. Homogenization temperatures were not determined, but the lack of retrograde reaction with garnet, such as the formation of epidote, suggests that the temperature of quartz veining is relatively high and beyond the range of epithermal-type mineralization.
There appears to be a transition from quartz veins with garnet envelopes to quartz veins and quartz flooding of the rock with no apparent reaction. Again, the lack of retrograde reaction with garnet, such as the formation of epidote, suggests that the temperature of quartz veining is relatively high. At the Campana mine, brown garnet skarn is cut by parallel quartz veins with a sheeted/ribbon texture. This rock clearly records two separate events. The first event is the formation of relatively coarse-grained garnet skarn with optical zonation similar to other Nambija samples (core composition Ad40, rim Ad60). Pyroxene in this sample has a similar iron content to the other Nambija samples, but the manganese content is even higher than the other samples (Hd31Jo13). The second event is a brittle deformation in which the rock has been veined by hundreds of parallel quartz veins. The walls of the quartz veins match perfectly, requiring that brittle fracture occurred without significant shear. The garnet crystals have been sliced, as if by a comb, into dozens of parallel slivers, with each sliver separated by optically continuous quartz. In both the quartz veins and quartz flooding there is no apparent reaction of the hydrothermal fluids with the wallrock (garnet). The fluid inclusions in the quartz are mostly vapor-rich indicating that boiling/fluid exsolution has occurred, probably due to a sudden pressure reduction (caused by fault movement?). There are no daughter minerals, so the total salinity is < 26 wt. % NaCl. This texture is similar to that observed in mesothermal orogenic gold deposits where quartz veins contain tens to hundreds of ribbons of sheared wall rock, separated by quartz.
Magnesian Gold Skarns
Although one million tons at an average grade of 6 g/t Au was produced from magnesian skarn at the Cable Mine, Montana (Earll 1972), most gold skarns are calcic skarns and little has been published until recently on the occurrence of magnesian gold skarns (Ettlinger et al. 1996; Mueller 1997). Most magnesian skarns form from dolomitic protoliths and exhibit a diagnostic mineralogy that includes forsterite, spinel, and serpentine. Although a variety of spinel phases can be present, magnetite usually is dominant and thus, most magnesian skarns are mined for iron and are relatively easy to find due to their strong magnetic signature. Butte Highlands, in southwest Montana is an unusual magnesian skarn in that it is an important gold resource, but lacks abundant iron oxides and sulfides. As pointed out by Ettlinger et al. (1996), the Fe-poor nature of this deposit means that it, and others like it, may not stand out during standard geophysical surveys.
Butte Highlands, Montana
Butte Highlands is one of many gold-bearing skarns associated with the relatively mafic marginal intrusions of the Boulder Batholith. Butte Highlands is located on the southern margin of the Boulder Batholith, about 24 km south of Butte, Montana. Skarn is associated with a fine- to medium-grained equigranular diorite, that has been intersected in drill core beneath the main mineralized area, called Nevin Hill (Ettlinger et al. 1996). Close to contacts with sedimentary rocks, the diorite exhibits endoskarn alteration with hornblende replaced by diopsidic pyroxene and titanite, plagioclase replaced by zoisite and prehnite, and calcium enrichment of plagioclase (to bytownite). In addition, the diorite is cut by veinlets of pyrrhotite with orthoclase, tremolite, and calcite envelopes.
The diorite has intruded the lower Paleozoic stratigraphic section at Butte Highlands causing extensive hornfelsing and recrystallization of the Wolsey, Meagher, Park, and Pilgrim formations. Argillaceous rocks of the Park Formation have been converted to biotite and pyroxene hornfels, similar to that described at many other gold skarns, whereas the dolomitic Meagher and Pilgrim formations have been recrystallized and locally silicified (Ettlinger et al. 1996). Mantos and chimneys of massive sulfide replacement ore in the Meagher and Pilgrim marbles were mined for base metals in the early part of the century (Sahinen 1950), but the bulk of skarn and gold mineralization occurs in the Wolsey Formation and the base of the Meagher Formation. The Wolsey Formation at Butte Highlands is described by Ettlinger et al. (1996) as consisting of interlayered, nonfossiliferous, dolomitic mudstone and shale, with some units of siltstone and carbonate.
Prograde skarn at Butte Highlands is dominated by forsteritic olivine with lesser pyroxene and phlogopite. Although these minerals are all pale green in color, this rock is black in hand specimen due to pervasive serpentization. Garnet is not abundant at Butte Highlands, but is present in endoskarn and with spinel as an overprint of the earlier olivine skarn. Such overprinting of early magnesian skarn minerals by later calcic skarn minerals has been reported from numerous magnesian skarn systems worldwide (Aksyuk & Zharikov 1988; Pertsev 1991). Neither the garnet nor spinel are Fe-rich, in contrast to most skarn systems. Retrograde alteration of olivine results in abundant serpentine, phlogopite, talc, carbonate, and magnetite. Retrograde alteration of more calcic skarn results in amphibole and vesuvianite, minerals that contain both Mg and Ca. Sulfide mineralization is strongly associated with retrograde alteration and Ettlinger et al. (1996) identified two associations with gold: phlogopite+pyrrhotite+gold and chlorite+clay+pyrrhotite+gold. In addition to this mineralogical association of gold with retrograde alteration, there is an elemental association of Au with Bi, based upon drill core assays (Ettlinger et al. 1996).
Skarn in "Mesothermal" Regional Metamorphic Terrains
Most skarns are associated with relatively shallow Phanerozoic plutons that have intruded previously unmetamorphosed sedimentary rocks (e.g., Einaudi et al. 1981). However, skarn mineralogy also has been described from several deposits in older orogenic belts where skarn is associated with both plutonism and high T-P metamorphism (e.g., Lucky Draw, Australia, Sheppard et al. 1995; Navachab, Namibia, Noertemann 1997; Tillicum, British Columbia, Ray et al. 1986a; Peterson 1996). In addition to these plutonic/metamorphic occurrences, there are several "mesothermal" lode gold deposits with skarn alteration in Precambrian terranes without associated intrusive rocks (e.g., Yilgarn craton, Western Australia, Mueller 1988, 1990, 1997, Mueller et al. 1991, 1996; Slave Province, northern Canada, Lhotka 1988, Lhotka & Nesbitt 1989, Bullis et al. 1994; Wyoming craton, USA, Smith 1996; Superior Province, eastern Canada, Hall & Rigg 1986, Pan & Fleet 1989, 1992, Pan et al. 1991; Dharwar craton, India, Siddaiah & Rajamani 1989).
These occurrences are significantly different from Phanerozoic skarn systems and little is known about the geologic relations of the skarn alteration or the connection between gold mineralization and skarn formation. Many researchers are unaware that these skarn occurrences even exist and there is much uncertainty about the timing and geochemistry of skarn formation. These skarns appear to be hybrids with characteristics of both the regional metamorphic environment and more typical Phanerozoic plutonism. What unites these disparate occurrences is a mineralogy dominated by very Fe-rich and reduced assemblages including garnet with major almandine-spessartine, hedenbergitic pyroxene, and Fe-rich amphibole. In some cases it appears that an Fe-rich protolith such as iron formation, komatiite, or metabasite is responsible for the unusual mineralogy. In addition, these deposits typically have part or all of the Au-As-Bi-Te geochemical signature of the younger gold skarn deposits. These "metamorphic" deposits are presented as a group because of their common link to regional metamorphism, even though there are huge differences in geologic setting and geochemistry among them. As more deposits like these are identified, it is hoped that understanding of their characteristics and origin will increase.
Lucky Draw, Australia
The Lucky Draw mine is located in the Burraga district 150 km west of Sydney in the Paleozoic Lachlan fold belt. Rocks in the Burraga district have been affected by two episodes of folding resulting in a series of upright, north-trending D1 anticlines and synclines. D2 folds are associated with Devonian-Carboniferous upper greenschist metamorphism and a regional slatey cleavage estimated to have formed at Ptotal = 2.0-2.5 kb (200-250 MPa) and T = 470° ± 35°C (Fowler 1987, 1989). Synchronous with regional metamorphism, a series of granitic plutons were emplaced with contact aureoles containing andalusite and cordierite. One of these, the Bathurst Granite, 30 km north of Lucky Draw, has been dated at 310 ± 7 Ma (Andrew 1984).
The Lucky Draw mine is located in the west-dipping limb of the D1 Brownlea anticline close to the contact with the Burraga Granodiorite. Mineralization occurs in the uppermost 100 m of the Ordovician Triangle Group, which consists of micaceous quartzite and quartz-mica schist, containing the assemblages quartz-albite-biotite-muscovite±cordierite and quartz-biotite-muscovite-albite-andalusite-cordierite, respectively (Sheppard et al. 1995). Within a few meters of the Burraga Granodiorite these assemblages are replaced by the assemblage quartz-biotite-plagioclase-cordierite-andalusite-sillimanite-Kfeldspar. Overlying the Triangle Group are the Rockley Volcanics, consisting of mafic and ultramafic flows, cumulates and volcaniclastic units. In the Burraga district, these rocks have been metamorphosed to tremolite-chlorite and quartz-feldspar-biotite-amphibole schists (Sheppard et al. 1995). Relict clinopyroxene and olivine phenocrysts compositions suggest that these schists were shoshonitic ultramafic cumulates and tuffs.
Alteration of Triangle Group quartzite and schist in the Lucky Draw mine area consists of an early metamorphic stage of medium- to coarse-grained gedrite, cordierite, and staurolite that define the metamorphic fabric of the rock, an intermediate stage of garnet-biotite-chlorite that veins and replaces earlier metamorphic minerals, and a late stage consisting of massive green-brown biotite spatially associated with the Burraga Granodiorite and small dikes. Temperatures of the early metamorphic stage are estimated by Sheppard et al. (1995) to be about 600°C based upon mineral equilibria. Temperature of the intermediate garnet-chlorite alteration was calculated to be 538 ± 62°C based upon Fe-Mg exchange between coexisting mineral pairs.
Mineralization at Lucky Draw is in the form of Au-As-Bi-Te minerals that are strongly associated with the garnet-chlorite stage of alteration. Identified minerals include ilmenite, arsenopyrite, molybdenite, native gold, maldonite (Au2Bi), native bismuth, bismuthinite, hedleyite (Bi14Te6), joseite-B (Bi4+xTe2-xS), tellurobismuthinite (Bi2Te3), and emplectite (CuBi2S) (Sheppard et al. 1995). The mineralization and all stages of alteration are very sulfide poor. Sheppard et al. (1995) state that pyrrhotite is the most abundant sulfide and estimate its abundance at less than 0.1% of mineralized sections.
The association of mineralization with the garnet-chlorite stage of metasomatic alteration suggests that introduction and/or remobilization of ore elements, Au-As-Bi-Te, occurred after the main phase of penetrative deformation and prior to the biotitization directly associated with crystallization and fluid exsolution from the Buragga Granodiorite. This is consistent with other studies, which have suggested that lode gold mineralization occurs post-peak metamorphism and is synchronous with or slightly earlier than plutonism (e.g., Mueller 1997). The calculated 538±62°C temperature of garnet-chlorite alteration is higher than the melting temperature of many of the ore minerals, suggesting that fluids circulating during garnet-chlorite alteration, perhaps driven by intrusion of the Buragga Granodiorite, leached ore elements from the adjacent mafic-ultramafic Rockley Volcanics and deposited them by reaction with the iron-rich minerals that are so abundant in the Lucky Draw mine area. Similar conclusions about leaching of mafic/ultramafic rocks during high temperature fluid circulation have been reached by other researchers (e.g., Steven 1993; Noertemann 1997).
Tillicum, British Columbia
Tillicum Mountain is located in south central British Columbia along the northern edge of the east trending Nemo Lakes Belt, a five kilometer wide roof pendant within the Cretaceous Nelson Batholith of upper greenschist- to lower amphibolite-grade metavolcanic and metasedimentary rocks, correlated with the Triassic Elise Formation of the Rossland Group (Peterson 1996). Regional metamorphism is sillimanite-grade at 5.0-6.8 kb (500-680 Mpa) and 630-680°C (Parrish 1981). In the Tillicum area, pressure and temperature conditions are thought to be slightly lower, 4.3-6.3 kb (430-630 Mpa) and 523-568°C , respectively and sillimanite does not occur (Ray et al. 1985; Peterson 1996). Metasedimentary rocks consist of thinly banded biotite-muscovite phyllite, spotted biotite schist, and graphitic biotite-muscovite phyllite. Metavolcanic rocks consist of shoshonitic and porphyritic mafic flows, tuffs, breccias, and intercalated argillites (Ray & Spence 1986).
In the vicinity of Tillicum Mountain, the Nelson Batholith consists of the Triassic Goatcanyon-Halifax Creek and mela-diorite stocks. The Goatcanyon-Halifax Creek stock is a medium-grained, equigranular quartz monzonite with an ilmenite to magnetite ratio of 5:1 (Peterson 1996). A marginal phase of the Goatcanyon-Halifax Creek stock is an equigranular mela-diorite or hornblendite containing xenoliths of the Goatcanyon-Halifax Creek stock. Based upon the hornblende geobarometer of Hollister et al. (1987), Peterson (1996) estimated pressures of emplacement of 6.3 and 4.3 kb, respectively, for the two plutons. Both stocks are undeformed and thus postdate skarn formation and mineralization associated with the main shearing event. In addition, the Goatcanyon-Halifax Creek stock truncates the trend of mineralization in underground workings (Peterson 1996).
Two episodes of folding have been documented in the Tillicum area (Ray et al. 1985; Peterson 1996). The first episode consists of southwest striking isoclinal folds and the development of a prominent axial planar schistosity (Peterson 1996). In a second episode, this schistosity was then refolded about a southwest plunging synform, accompanied by shearing. Skarn formation and mineralization can be placed within this structural framework, because calc-silicate minerals overgrow and cut the D1 metamorphic fabric in the district. In detail, native gold, native bismuth, and bismuthinite occur in fractures in garnet, gold occurs along cleavage planes in pyroxene, and massive sulfide locally replaces calc-silicate skarn. In addition, skarn minerals and mineralized veins are locally sheared and folded by D2. Thus, skarn formation and mineralization is post D1 and roughly synchronous with D2. On a local scale, Peterson (1996) suggested that deformation was concentrated along the margins of metavolcanic flows and sills that behaved as competent blocks within a weaker metasedimentary matrix and that hydrothermal/metamorphic fluids were focused along these contacts.
All garnets associated with mineralization are subcalcic. A small calcareous unit in the North Slope area close to the Goatcanyon-Halifax Creek intrusion has grandite garnets, but this occurrence is not mineralized. Metamorphic garnets and skarn garnets related to mineralization have similar but distinct compositions. Metamorphic garnets occur in banded schists and gneisses as part of the D1 penetrative deformation. They are more subcalcic than skarn garnets, containing as little as 6 mole % grandite in contrast to a maximum of 64 mole % grandite in skarn garnets. Moreover, on average they contain twice as much pyrope (up to 22 mole %) and half as much spessartine (as low as 15 mole %) as do skarn garnets (as little as 1 mole % pyrope and as much as 61 mole % spessartine). Both pyroxene and amphibole are the typical calcic varieties that occur in "normal" skarn deposits in a non-regional metamorphic environment. Pyroxene is diopsidic (Hd4-42) and amphibole is mostly within the tremolite-actinolite series (Peterson 1996).
Using the garnet-biotite geothermometer of Ferry & Spear (1978), Peterson (1996) calculated a temperature of 523-568°C for garnet skarn formation, assuming a pressure of 6 kb (600 MPa) as determined for the later emplacement of the Goatcanyon-Halifax Creek stock. This temperature range is consistent with the fluid inclusion trapping temperatures (500-550°C) determined from garnet and quartz associated with mineralization, using a pressure correction based upon the aforementioned pressure of 6 kb (600 MPa). No daughter minerals were observed in any fluid inclusions and fluid inclusions in quartz yielded salinities of 0.7-3.8 eq. wt. % NaCl (Peterson 1996).
The other major alteration type in the Tillicum area is biotite hornfels, which affects all the rock types in the district, except the Goatcanyon-Halifax Creek and mela-diorite intrusions. Depending on the protolith being altered, biotite hornfels can be texturally diverse but always contains biotite, quartz, and K-feldspar and usually is unfoliated. More felsic protoliths tend to have more K-feldspar and more mafic protoliths, more biotite. Some of the biotite alteration of metavolcanic rocks, such as the diorite porphyry flows and sills, is relatively coarse grained with biotite up to several mm. Thus, the term biotite hornfels is not strictly appropriate, but has been retained as a field term because biotite is an essential component and the alteration typically is fine grained and granular.
There are six mineralized zones at Tillicum: Heino-Money, East Ridge, Silver Queen, North Slope, Grizzly, and Arnie Flats. Reserves at East Ridge are 1.4 Mt averaging 7.7 g/t Au. Heino-Money is smaller but much higher grade, with 55,000 tons averaging 33.4 g/t Au. Mineralization is spatially associated with skarn alteration and shear zones. Some of the shearing predates skarn, as the calc-silicate minerals overgrow the penetrative fabric of the earlier metamorphism. In other cases, sheared quartz veins with coarse-grained visible gold have skarn envelopes and some of the calc-silicate minerals are weakly foliated. Thus, it appears that the Tillicum area has been structurally active for a considerable time and that for at least part of that time, skarn-forming hydrothermal fluids were active during shearing. Veins consist of quartz, calcite, pyroxene, amphibole, clinozoisite, garnet, Kfeldspar, titanite, biotite, and muscovite. Sulfide minerals include major pyrrhotite and pyrite. Minor to trace minerals include native gold, marcasite, native bismuth, bismuthinite, hedleyite, and joesite-B (Peterson 1996).
In general, mineralization in the Tillicum area is not sulfide-rich. One exception is in the Heino-Money zone where pyroxene-amphibole-calcite skarn is replaced by a vein of massive sulfide consisting of pyrrhotite, sphalerite, galena, boulangerite, arsenopyrite, chalcopyrite, and freibergite. Sulfide replacement ranges from 20 to 95% of the rock, averaging 80%. Values of Pb, Zn, and Ag range up to 7.2%, 39.5%, and 100 oz/t, respectively, with gold values of 0.2 to 1.0 oz/t. Massive sulfide replacement zones also occur in the East Ridge zone and although they are geochemically anomalous in Au, Ag, As, and base metals, they are not ore grade (Peterson 1996).
The open pit mine at Navachab is located 10 km south of Karibib in the southern central zone of the Damara Orogen (Pirajno & Jacob 1991; Moore and Jacob, 1998). Production is 1.8-1.9 t/y Au from a reserve of 9.75 Mt at an average grade of 2.4 g/t Au from pyroxene-rich zones in skarn formed in metasedimentary rocks of the Okawayo Formation of the Swakop Group (Noertemann 1997). Regionally, the Swakop Group includes the Spes Bona Formation, consisting of schists, calc-silicate rocks, and meta-arkoses, the Karibib Formation, consisting of a basal calc-silicate white marble, a dolomitic brown marble, and a hanging wall gray graphitic marble, the Okawayo Formation consisting of calc-silicate marbles, and the Oberwasser Formation, consisting of siliciclastic units (Steven 1993). Within the Okawayo Formation is a distinctive dark rock, which in the mine is called the marker hornfels, but which geochemically is a metamorphosed, late Damaran, camptonitic lamprophyre (Noertemann 1997). This rock served as a fluid barrier and chemical trap for mineralizing fluids. Mineralized skarn is immediately adjacent to the meta-lamprophyre, but the lamprophyre itself is barren of gold (Noertemann 1997).
Puhan (1983) documented metamorphic P-T-conditions ranging from 2.6-3.4 kb (260-340 Mpa) and 555°C-645°C in the central Damara Orogen. Noertemann (1997) showed that the area has been affected by a combination of polyphase folding and late-tectonic brittle-ductile shearing. The D1 deformation produced East-verging recumbent F1 folds on a scale of several kilometers. Subsequently, this folding was overprinted by a progressive F2 folding, as indicated by refolding and intrafolial folds. After that, an isoclinal F3 folding led to the formation of a weakly NW-verging, upright D3 anticline on a regional scale, which hosts several gold deposits in the south central Damara Orogen, including Navachab.
Navachab represents a reduced distal and Mn-enriched gold skarn formed in banded, predominantly calcite marble with biotite schist and calc-silicate layers. Regional metamorphism of these rocks produced preferred growth of garnet in pelitic layers and of clinopyroxene in carbonate layers. These early metamorphic garnets are intermediate grandite-pyralspite and the metamorphic pyroxenes are salite with only minor johansenite. Metasomatic skarn, veins and overprints these layered metamorphic occurrences and both garnet and pyroxene are enriched in manganese. This time progression is clearly visible in garnets which have a metamorphic poikiloblastic core and a younger, inclusion-free margin. Associated amphibole is largely tremolite-actinolite with a minor hastingsite component. Amphibole shows a strong decrease in Mg and increase in Fe2+, Mn2+, and Fe3+ from marble to skarn. The occurrence of graphite in the skarn as a product of decarbonatisation implies very low oxygen-fugacities, consistent with the lack of magnetite and hematite (Noertemann 1997).
The ore mineralization is distinguished into two main parageneses: pyrrhotite, chalcopyrite, arsenopyrite, molybdenite and sphalerite associated with early skarn formation and a younger one with remobilised pyrrhotite, chalcopyrite, pyrite, native bismuth, bismuthinite, maldonite, and native gold, which is associated with skarn and retrograde amphibole alteration. Noertemann (1997) estimated the P-T conditions of ore formation by geobarometry and geothermometry of sphalerite [2-2.5 kb (200-250 Mpa) and 590 °C] and arsenopyrite (575 ± 15°C). As in other regional metamorphic skarn occurrences, these temperatures and pressures are slightly lower than the peak conditions as determined by Puhan (1983).
Lupin, Northwest Territories, Canada
The Lupin deposit is located 400 km NNE of Yellowknife in the Contwoyto Lake-Point Lake area of the Archean Slave Province of the Canadian Shield, and represents the largest (11.8 Mt at an average grade of 10.0 g/t Au, for a total of 117 t Au) of more than 100 gold occurrences in this area (Bullis et al. 1994). The Lupin deposit consists of a series of stratabound replacement orebodies developed in iron formation adjacent to cross-cutting quartz veins. The complexly folded iron formation is intercalated with Archean greywackes and turbidites of the Contwoyto Formation. The deposit occurs within the broad metamorphic aureole of a large granodiorite-granite pluton, the Contwoyto batholith, at a distance of about 1.5 km to the south of the contact. In the mine area, the unmineralized iron formation consists of mesobanded quartz and grunerite and is metamorphosed to upper greenschist facies grade (Lhotka 1988). The cordierite isograd, marking the position of amphibolite-facies metamorphism, passes about 400 m to the north of the mine at surface, but intersects the mine workings at a depth of 550 m (Lhotka & Nesbitt 1989). Garnet-biotite pairs from the Lupin mine indicate a temperature of 600°C at 3 kb (300 MPa), slightly higher than the 575°C determined from cordierite mineral assemblages (Lhotka 1988).
The region has been effected by at least three deformation events (King et al. 1988; Relf 1989). The first developed prior to the peak of regional metamorphism and consists of tight isoclinal folds in which the S1 axial planar cleavage is defined by the alignment of biotite and muscovite. The second deformation phase developed during peak metamorphic conditions and F2 folds are tight to isoclinal with steep plunges and near-vertical axial planes. F3 folding created crenulations to earlier folds. The Contwoyto batholith is thought to have intruded during D3. Quartz veins are abundant in the mine and appear to be localized in fold hinges and may be related to the Lupin Fault, which bounds the orebodies to the southwest (Lhotka 1988). Although not ore grade, the quartz veins locally contain native gold, pyrrhotite, arsenopyrite, and scheelite (Lhotka 1988).
The main ore host at Lupin is an iron formation that has been metamorphosed to an assemblage of grunerite-quartz-magnetite and then later retrograded/sulfidized to include hornblende, ilmenite, and pyrrhotite. In sulfide-rich iron formation, particularly near quartz veins, almost all the grunerite is replaced by hornblende, and arsenopyrite, loellingite, and pyrite are present in addition to pyrrhotite. In zones of very intense alteration/sulfidation, calcic garnet, pyroxene, and actinolite are also present (Lhotka 1988). Locally, garnet amphibolite occurs as lenses within or along the margins of iron formation. The garnet amphibolite consists of almandine garnet-grunerite-chlorite and contains hedenbergitic pyroxene near contacts with iron formation (Bullis et al. 1994). Retrograde alteration of these rocks, again associated with quartz veins, includes epidote and actinolite in addition to hornblende.
Lupin contains gold mineralization associated with calc-silicate alteration that is similar to many Phanerozoic gold skarns. In contrast to some of the Phanerozoic gold skarns associated with regional metamorphism where the protoliths typically contain at least some calcium to form calcic garnet and pyroxene, the iron formation host at Lupin is very calcium poor and Lhotka (1988) determined that Ca was introduced metasomatically by hydrothermal fluids. The source of the calcium and the ultimate source of the hydrothermal fluids are not known. However, intrusive dikes of felsic to intermediate composition, common in many Phanerozoic skarn deposits, are exposed on the lower levels of the mine. The petrology of these dikes and their relationship to the gold skarn orebodies are not known (Lhotka & Nesbitt 1989).
Nevoria, Western Australia
The Nevoria, Marvel Loch, Big Bell, and other gold skarn deposits in Western Australia are located in the mid-central part of the Archean Yilgarn Craton (Mueller 1997). The skarns are hosted by ultramafic to mafic volcanic rocks or banded iron-formation of the Southern Cross Belt, a narrow greenstone terrane of greenschist to amphibolite grade. They are controlled by NW-trending ductile shear zones and occur within the broad, contact-metamorphic aureole of the Ghooli Dome, a composite granodiorite-granite batholith located to the southwest of a major transcurrent shear zone of more than 300 km strike length. Pressure and temperature estimates for peak metamorphism [4 ± 1 kb (400 Mpa) and 570-610°C], suggest that the present erosion surface in the Southern Cross area and the gold deposits it exposes were at a depth of 10 to 15 km during contact metamorphism and subsequent skarn formation (Mueller 1997).
The region has been affected by multiple episodes of deformation. Symmetrical repetition of units on both sides of the Ghooli Dome suggests an early regional fold with steeply dipping (70-80°S) limbs. This fold is then refolded by the second episode of ductile deformation with north-south-striking axial planes and lineations that plunge 70-80°E. Both generations of folds are cut by brittle-ductile faults that strike N10-25°E and dip 55-65°E; fault striations plunge 10-35°S. These faults cut skarn alteration (Mueller 1997).
The Nevoria gold skarn represents a smaller (0.4 Mt underground, averaging 6.3 g/t Au and 3.15 Mt open pit, averaging 2.7 g/t Au) analogue of the Lupin deposit in Canada (Mueller 1997). The Nevoria deposit is hosted by silicate banded iron-formation within amphibolite-facies greenstones overlying the Ghooli Dome batholith. Deep drill holes have intersected massive pegmatite and biotite granite of the Ghooli Dome at a vertical depth of 250 m below the Nevoria open pits, and post-mineralization pegmatite dikes cut across the orebodies at the higher levels of the mine (Mueller 1988). The orebodies occur in three steeply dipping horizons of grunerite-quartz iron formation, interbedded with fine-grained ultramafic to mafic amphibolites. Skarn is present in both iron formation and in the ultramafic to mafic volcanic rocks, but gold mineralization is usually of subeconomic grade in the latter (Mueller 1990). Even in the iron formation, the distribution of gold is not homogeneous because high-grade skarns are localized near contacts with metavolcanic rocks. The central portions of the iron formations are less altered and poorly mineralized.
The skarn orebodies at Nevoria occur as stratabound replacements of grunerite-quartz iron formation adjacent to flatly dipping quartz veins. The major prograde alteration minerals are coarse-grained hedenbergite, locally intergrown with minor calcite and rare scheelite, and medium-grained ferro-actinolite or ferro-hornblende. Minor almandine garnet with ferro-hornblende, biotite, ilmenite, and tourmaline occurs in bands parallel to the iron formation mesobanding, and at the margins of cross-cutting quartz veins. Grandite garnet-pyroxene-actinolitic hornblende veins occur in tholeitic and komatiitic amphibolites, but not in iron formation. Retrograde alteration minerals include Fe-rich chlorite, muscovite, and stilpnomelane. The gangue minerals are intimately intergrown with disseminated to massive pyrrhotite, and minor chalcopyrite, loellingite, arsenopyrite, pyrite, and scheelite. Gold occurs in native form, either enclosed in pyrrhotite or in skarn silicates (e.g., hedenbergite; see Fig. 10c in Mueller 1997), and is often associated with tsumoite (BiTe), Bi-Te sulfides (e.g., Bi3.11Te0.85S1.04), and maldonite (Au2Bi), which locally has exsolved to native bismuth and native gold (Mueller 1997).
Skarn is also present in the ultramafic to mafic amphibolites intercalated with the iron formation horizons, but is subeconomic in grade. The most prominent alteration features in the more massive amphibolites are zoned garnet-pyroxene replacement veins with highly irregular boundaries (Mueller 1988). The veins consist of a core of grossular garnet and outer margins of diopside. Minor plagioclase, microcline and scheelite are intergrown with both garnet and pyroxene. The peak fluid temperature during the formation of the zoned skarn veins in the ultramafic to mafic volcanic rocks, is constrained by the reaction clinozoisite + quartz + calcite = grossular to values of 550-580°C, assuming a pressure of 4 kb (400 MPa) and low mole fraction CO2 (0.03-0.05) in the fluid. Retrograde minerals in the veins include the assemblage clinozoisite + calcite + quartz, filling cracks in grossular garnet, and aggregates of muscovite and prehnite replacing feldspar. Disseminated sulfides are rare, and consist mainly of pyrite and chalcopyrite (Mueller 1990).
The difference between skarn hosted in iron formation and in amphibolite goes beyond mineralogy. Mueller (1997) postulated that iron formation and amphibolite have contrasting physical properties and that, during D1 folding, contacts between these two rock types accommodated most of the differential slip. These contacts were then the locus of preferential alteration and mineralization.
In conclusion, gold skarns occur worldwide and in a variety of geologic settings. These deposits share many common features such as biotite hornfels, garnet-pyroxene alteration, clastic- and/or volcaniclastic-rich protoliths, and a Au-As-Bi-Te geochemical signature, but also exhibit significant differences, especially among the four major subdivisions described in this review: 1) reduced Au skarns, 2) oxidized Au skarns, 3) magnesian Au skarns, and 4) metamorphic Au skarns. Both reduced and oxidized Au skarns are related to shallow Phanerozoic plutons. Most depth estimates for these systems are < 5 km, broadly similar to the general environment of porphyry-type deposits. Plutons associated with reduced gold skarns tend to be ilmenite-bearing mafic diorites and granodiorites, whereas plutons associated with oxidized gold skarns tend to be more silicic and magnetite-bearing (Meinert 1995; Ray et al. 1995). In contrast, magnesian and metamorphic Au skarns do not necessarily occur with associated igneous rocks and they range in age from Archean to Phanerozoic. The formation of skarn in these systems appears to be more dependent on particular host rock compositions and relatively high P-T conditions rather than on the petrochemistry of associated plutons. Yet even with these fundamental differences, most of the deposits still exhibit biotite hornfels, garnet-pyroxene alteration, and a Au-As-Bi-Te geochemical signature. Thus, it seems that fundamentally similar geochemical processes can occur in what appear to be relatively disparate geological environments. A similar conclusion has been reached for other geologic environments, such as geothermal (Bird et al. 1984) and VMS (Galley & Ames 1998) deposits, some of which also exhibit skarn alteration.
The biotite±Kfeldspar (potassic) alteration that surrounds most gold skarns is one of the characteristics of this deposit type. In most cases, the biotite±Kfeldspar alteration forms in relatively fine-grained, clastic host rocks, resulting in a hornfels texture. However, in coarser-grained rocks, particularly igneous rocks, the biotite±Kfeldspar alteration can be relatively coarse-grained, as has been described at Hedley in the reduced skarn class (Ettlinger et al. 1992; Ray et al. 1996b) and Tillicum in the metamorphic skarn class (Peterson 1996). On the scale of a hand specimen, the biotite±Kfeldspar alteration clearly is metasomatic both in the sense of fluid infiltration via fractures and in the sense of mass transfer, e.g., K is being added to the rock. However, on the larger deposit scale, much of the biotite±Kfeldspar alteration appears to be more a case of in situ redistribution during progressive mineral reactions, as documented at Beal (Wilkie 1996). In most cases, the original clastic sedimentary rocks contain abundant K-rich argillaceous and arkosic material which has been replaced by pyroxene and then garnet in inner or more intense zones of alteration. Since garnet and pyroxene do not contain K, the K originally present in the host rocks is liberated during skarn formation and may be incorporated into the distal biotite±Kfeldspar alteration. In contrast to other elements such as Fe and Au, which are demonstrably added to the alteration assemblage from magmatically derived hydrothermal fluids (e.g., Zimmerman et al. 1992), K in biotite±Kfeldspar alteration appears to be, at least in part, internally derived from the host stratigraphy. This conclusion is supported by the fact that skarns developed in relatively pure limestone do not generally develop biotite hornfels or other K-rich alteration.
If biotite±Kfeldspar hornfels is thought of as the distal alteration zone of gold skarns, then the proximal part is relatively coarse-grained garnet and pyroxene. As documented for numerous deposits, this proximal skarn zone is internally zoned from garnet-dominant close to the pluton or fluid pathway to pyroxene-dominant away from the pluton or fluid pathway (Meinert 1997). The relative proportion of garnet and pyroxene is a complex function of protolith composition, activity of components in the hydrothermal fluid, and overall oxidation state as influenced by magmatic sources, wall-rock composition, and mineral reactions (e.g., Newberry 1991). In a general way this corresponds to the fundamental classification of a given skarn as reduced versus oxidized. The oxidation state of a skarn system is reflected in both the relative proportions of garnet and pyroxene and in the composition of individual mineral phases.
The most reduced skarn assemblages lack ferric iron and contain major amounts of ferrous iron in both garnet and pyroxene. Generally, such compositions only occur in skarns, such as Lupin, Navachab, and Nevoria, formed from Fe-rich rocks at great depth in regional metamorphic terrains. Pressures in these deposits have been estimated at 2-6 kb (200-600 MPa), in contrast to < 0.5 kb (50 MPa) in the less reduced skarns associated with hypabyssal intrusions. W skarns provide a link between these two groups in that some of the deeper W skarns with particularly reduced graphitic wall rocks have hedenbergitic pyroxene and Fe-rich subcalcic garnets (Newberry 1983). Although W skarn garnets do not contain nearly as much almandine component as do the garnets in metamorphic Au skarns, the subcalcic almandine component can be taken as an indirect measure of very low oxidation states. The next most reduced mineral assemblage lacks significant ferric iron in garnet, has major amounts of ferrous iron in pyroxene, but lacks ferrous iron in garnet. Au skarns with such hedenbergitic pyroxene and intermediate grandite garnet include Crown Jewel, Elkhorn, Fortitude, Hedley, and Junction Reefs. The least reduced mineral assemblage is characterized by major amounts of andraditic garnet and diopsidic to salitic pyroxene. This assemblage is characteristic of all the skarns in the oxidized Au skarn category.
Although most sulfide mineralization occurs at lower temperatures after the main stage of garnet and pyroxene formation, the silicate, oxide, and sulfide mineral association of a given skarn rock tends to reflect the overall oxidation and sulfidation state of the hydrothermal system. This can be represented for Au skarn deposits by considering relevant silicate, oxide, and sulfide mineral equilibria. Oxidized Au skarns typically contain andraditic garnet, diopsidic pyroxene, tremolite-actinolite, pyrite, and hematite. Reduced Au skarns typically contain intermediate grandite garnet, hedenbergitic pyroxene, actinolite-ferro-actinolite, pyrrhotite, ± magnetite. Metamorphic Au skarns typically contain subcalcic garnet with significant almandine-spessartine, hedenbergitic pyroxene, grunerite-ferro-actinolite, pyrrhotite, and graphite.
The metamorphic Au skarn category provides an important link to gold deposits, typically described as orogenic or mesothermal lode Au deposits, that are associated with shear zones in deep crustal environments (e.g., Mueller & Groves 1991; McCuaig & Kerrich 1994). These deposits are characterized by sericite-carbonate-albite wall rock alteration and lack calc-silicate minerals. Even where the protoliths are particularly Fe-rich, such as iron formation, calc-silicate minerals are lacking (e.g., Williams 1997). The metamorphic Au skarns form in the same general geologic environment and under similar P-T conditions. They differ mainly in the composition of the fluid phase and the resulting metasomatism that stabilizes calc-silicate minerals. Mueller & Groves (1991) suggested that skarn-forming metasomatism in shear zones is restricted to deeper, higher temperature occurrences. Other workers have ascribed this metasomatic signature to a particular igneous petrogenesis (e.g., Spooner 1993) but the apparent lack of magmatism in some districts like Navachab suggests that co-genetic magmatism is not required in all cases.
The Au-As-Bi-Te geochemical signature of Au skarns also can be related to geochemical variations involving temperature, f(O2), and f(S2). For example there are two main As minerals in gold skarns, arsenopyrite, and loellingite. The assemblage loellingite-pyrrhotite is only stable at a combination of relatively high temperature and relatively low sulfidation state. Such conditions are typical of the very reduced metamorphic gold skarns and loellingite-pyrrhotite is reported from all these deposits. In many deposits, such as Navachab (Noertemann 1997), loellingite is described as rimmed or partially replaced by arsenopyrite, particularly where associated with retrograde alteration of associated silicate minerals. This is consistent with isobaric cooling at low f(O2) and f(S2) conditions. In the reduced Au skarns associated with Phanerozoic hyabyssal intrusions, loellingite is less common than arsenopyrite in all deposits and is absent from some. Arsenopyrite in these deposits typically occurs with pyrrhotite and generally is reported to be relatively late in the paragenesis, e.g., at relatively low temperature compared to the early garnet and pyroxene. In the oxidized Au skarns, loellingite is absent and arsenopyrite is not abundant, typically occurring with pyrite and as late cross cutting veins, indicative of relatively high sulfidation states and/or low temperatures.
One of the apparent anomalies of Au skarn mineralogy is the ubiquitous occurrence of Bi, Bi-Au, and Bi-Te minerals with very low melting points, in most cases much lower than the formation temperature of associated silicate, oxide, and sulfide minerals. For example, maldonite (Au2Bi) unmixes into native Au and native Bi at a temperature of 113°C (Barton & Skinner 1979). Most of the Bi, Bi-Au, and Bi-Te minerals in Au skarns probably were intermediate composition solid solutions at the temperature of initial skarn formation and their present mineralogy represents a long history of unmixing and remobilization. Particularly in metamorphic environments, it might be predicted that these elements will have a long and varied history during protracted thermal, hydrothermal, and structural events.
In contrast to many other types of gold deposits, such as the epithermal class in which transport of gold as bisulfide complexes at temperatures < 300°C and the mechanical/chemical effects of boiling are critical features, Au skarns contain abundant evidence for high-temperature transport of Au in saline to hypersaline fluids. At the high temperatures documented for many skarn deposits (350-650°C), Au could be transported in high concentrations as chloride complexes (Gammons & Williams-Jones 1995, 1997), whereas at temperatures below 300°C, a thiosulfide complex [Au(HS)2-] is more likely (Hayashi & Ohmoto 1991). One of the objections to chloride transport of Au in even relatively high-temperature mesothermal Au deposits is the lack of base metals in these deposits even though Fe, Cu, Zn, etc. are highly soluble in high-temperature chloride fluids (McCuaig & Kerrich 1994). This objection does not hold for most Au skarns as base metals, particularly Fe, are common.
The problem then becomes more one of getting gold out of chloride solutions rather than into solution. For Au skarns associated with plutons, the inevitable cooling of hydrothermal fluids away from the magma and as the magma itself cools, provides a ready mechanism for depositing gold due to the relatively low solubility of Au chloride complexes below 400°C, and especially below 300°C. Meinert (1989) suggested coupled transport for such Au skarns, with chloride transport at high temperatures and bisulfide transport becoming increasingly important below about 350°C. Due to the broadly isothermal conditions during regional metamorphism, this mechanism is less likely for the regional metamorphic Au skarns. In these systems, the mineralogical evidence for fluid-wallrock reaction suggests that Au deposition may be due to sulfidation of Fe-rich wallrocks such as at Lupin (Bullis et al. 1994) or pH changes due to K metasomatism or reaction with carbonate wallrocks. Alternatively, the relatively low oxidation and sulfidation states of most Au skarn deposits provide a clue that reduction of an oxidized hydrothermal fluid may be a fundamental cause for the precipitation of gold in all of these systems.
Depending on whether one is a "lumper" or a "splitter", Au skarn deposits are either fairly similar since they all contain calc-silicate minerals or they are vastly different because they span such a wide variety of host rocks, intrusive relationships, and details of mineral chemistry. What unites them is a common mineralogy and geochemistry rooted in the definition of skarn. But it is not the individual minerals that provide the key to understanding these fascinating deposits, but rather the underlying geological and chemical processes that control mineral stabilities and control whether a system has the potential for turning into a profitable mine or remaining a mineralogical curiosity. Although relatively young Phanerozoic gold skarns directly associated with shallow intrusions are the best known of the systems described in this review, the presence of skarn-type systems in older rocks and associated with the much larger scale of regional metamorphism, suggests a new target of opportunity for exploration and research. Just as in the case when serious exploration of gold skarn systems was spurred by the dramatic rise in the price of gold during the 1980s, the re-examination of old familiar metamorphic terrains with a fresh eye, a fresh "skarn eye", may yield significant new discoveries and yet another new gold rush for the next century.
Tungsten skarns are found on most continents in association with calc-alkaline plutons in major orogenic belts. Major reviews of tungsten skarns include Newberry and Einaudi (1981), Newberry and Swanson (1986), Kwak (1987), and Newberry (1998). As a group, tungsten skarns are associated with coarse-grained, equigranular batholiths (with pegmatite and aplite dikes) surrounded by large, high-temperature, metamorphic aureoles. These features are collectively indicative of a deep environment. Plutons are typically fresh with only minor myrmekite and plagioclase-pyroxene endoskarn zones near contacts.
The high-temperature metamorphic aureoles common in the tungsten skarn environment contain abundant calc-silicate hornfels, reaction skarns, and skarnoid formed from mixed carbonate-pelite sequences. Such metamorphic calc-silicate minerals reflect the composition and texture of the protolith and can be distinguished from ore-grade metasomatic skarn in the field and in the laboratory.
Newberry and Einaudi (1981) divided tungsten skarns into two groups: reduced and oxidized types, based on host rock composition (carbonaceous versus hematitic), skarn mineralogy (ferrous versus ferric iron), and relative depth (metamorphic temperature and involvement of oxygenated groundwater). Early skarn assemblages in reduced tungsten skarns are dominated by hedenbergitic pyroxene and lesser grandite garnet with associated disseminated fine-grained, molybdenum-rich scheelite (powellite). Later garnets are subcalcic (Newberry, 1983) with significant amounts (up to 80 mole %) of spessartine and almandine. This subcalcic garnet is associated with leaching of early disseminated scheelite and redeposition as coarse-grained, often vein-controlled, low-molybdenum scheelite. It is also associated with the introduction of sulphides, such as pyrrhotite, molybdenite, chalcopyrite, sphalerite, and arsenopyrite, and hydrous minerals such as biotite, hornblende, and epidote.
In oxidized tungsten skarns, andraditic garnet is more abundant than pyroxene, scheelite is molybdenum-poor, and ferric iron phases are more common than ferrous phases. For example at the Springer deposit in Nevada, garnet is abundant and has andraditic rims, pyroxene is diopsidic (Hd0-40), epidote is the dominant hydrous mineral, pyrite is more common than pyrrhotite, and subcalcic garnet is rare to absent (Johnson and Keith, 1991). In general, oxidized tungsten skarns tend to be smaller than reduced tungsten skarns, although the highest grades in both systems typically are associated with hydrous minerals and retrograde alteration.
Copper skarns are perhaps the worlds most abundant skarn type. They are particularly common in orogenic zones related to subduction, both in oceanic and continental settings. Major reviews of copper skarns include Einaudi et al. (1981) and Einaudi (1982a,b). Most copper skarns are associated with I-type, magnetite series, calc-alkaline, porphyritic plutons, many of which have co-genetic volcanic rocks, stockwork veining, brittle fracturing and brecciation, and intense hydrothermal alteration. These are all features indicative of a relatively shallow environment of formation. Most copper skarns form in close proximity to stock contacts with a relatively oxidized skarn mineralogy dominated by andraditic garnet. Other phases include diopsidic pyroxene, idocrase, wollastonite, actinolite, and epidote. Hematite and magnetite are common in most deposits and the presence of dolomitic wall rocks is coincident with massive magnetite lodes which may be mined on a local scale for iron. As noted by Einaudi et al. (1981), copper skarns commonly are zoned with massive garnetite near the pluton and increasing pyroxene and finally idocrase and/or wollastonite near the marble contact. In addition, garnet may be color zoned from proximal dark reddish-brown to distal green and yellow varieties. sulfide mineralogy and metal ratios may also be systematically zoned relative to the causative pluton. In general, pyrite and chalcopyrite are most abundant near the pluton with increasing chalcopyrite and finally bornite in wollastonite zones near the marble contact. In copper skarns containing monticellite (e.g. Ertsberg, Irian Jaya, Indonesia, Kyle et al., 1991; Maid of Erin, British Columbia, Meinert unpub. data) bornite-chalcocite are the dominant Cu-Fe sulfides rather than pyrite-chalcopyrite (e.g. Big Gossan, Irian Jaya, Meinert et al., 1997). The largest copper skarns are associated with mineralized porphyry copper plutons. These deposits can exceed 1 billion tons of combined porphyry and skarn ore with more than 5 million tons of copper recoverable from skarn. The mineralized plutons exhibit characteristic potassium silicate and sericitic alteration which can be correlated with prograde garnet-pyroxene and retrograde epidote-actinolite, respectively, in the skarn. Intense retrograde alteration is common in copper skarns and in some porphyry-related deposits may destroy most of the prograde garnet and pyroxene (e.g. Ely, Nevada; James, 1976).
Endoskarn alteration of mineralized plutons is rare. In contrast, barren stocks associated with copper skarns contain abundant epidote-actinolite- chlorite endoskarn and less intense retrograde alteration of skarn. Some copper deposits have coarse-grained actinolite-chalcopyrite-pyrite-magnetite ores but contain only sparse prograde garnet-pyroxene skarn (e.g. Monterrosas and Ral-Condestable deposits, Peru: Ripley and Ohmoto, 1977; Sidder, 1984; Vidal et al., 1990; Record mine, Oregon, Caffrey, 1982; Cerro de Mercado, Mexico, Lyons, 1988). These deposits provide a link between some copper and iron skarns and deposits with volcanogenic and orthomagmatic affinities.
Most zinc skarns occur in continental settings associated with either subduction or rifting. They are mined for ores of zinc, lead, and silver although zinc is usually dominant. They are also high grade (10-20% Zn+ Pb, 30-300 g/t Ag). Related igneous rocks span a wide range of compositions from diorite through high-silica granite. They also span diverse geological environments from deep-seated batholiths to shallow dike-sill complexes to surface volcanic extrusions. The common thread linking most zinc skarn ores is their occurrence distal to associated igneous rocks. Major reviews of zinc skarn deposits include Einaudi et al. (1981) and Megaw et al. (1988).
Zinc skarns can be subdivided according to several criteria including distance from magmatic source, temperature of formation, relative proportion of skarn and sulfide minerals, and geometric shape of the ore body. None of these criteria are entirely satisfactory because a magmatic source cannot be identified for some deposits, because most skarns develop over a range of temperatures, and because most large skarn deposits contain both skarn-rich ores and skarn-poor ores in a variety of geometric settings including mantos and chimneys (e.g. Megaw, 1998). Megaw et al. (1988) make the important point that many zinc skarn districts "grade outward from intrusion-associated mineralization to intrusion-free ores, which suggests that those districts lacking known intrusive relationships may not have been traced to their ends". Similarly, most zinc skarn districts grade outward from skarn-rich mineralization to skarn-poor ores, veins, and massive sulfide bodies which may contain few if any skarn minerals. Incompletely explored districts may only have some of these zones exposed. But as previously noted, the presence of skarn minerals, such as garnet and pyroxene within the system, is important because it indicates a restricted geochemical environment which is entirely distinct from ore types, such as Mississippi Valley-type deposits, which also contain Zn-Pb-Ag ores but which absolutely lack skarn minerals.
Besides their Zn-Pb-Ag metal content, zinc skarns can be distinguished from other skarn types by their distinctive manganese- and iron-rich mineralogy, by their occurrence along structural and lithologic contacts, and by the absence of significant metamorphic aureoles centered on the skarn. Almost all skarn minerals in these deposits can be enriched in manganese including garnet, pyroxene, olivine, ilvaite, pyroxenoid, amphibole, chlorite, and serpentine.
In some deposits, the pyroxene:garnet ratio and the manganese content of pyroxene increase systematically along the fluid flow path (e.g. Groundhog, New Mexico, Meinert, 1987). This feature has been used to identify proximal and distal skarns and proximal and distal zones within individual skarn deposits. A typical zonation sequence from proximal to distal is: altered/endoskarned pluton, garnet, pyroxene, pyroxenoid, and sulfide/oxide replacement bodies (sometimes called mantos and chimneys based upon geometry and local custom). The occurrence of zinc skarns in distal portions of major magmatic/hydrothermal systems may make even small deposits potentially useful as exploration guides in poorly exposed districts. Thus, reports of manganese-rich mineral occurrences may provide clues to districts that have not yet received significant exploration activity. Another evidence of distal hydrothermal alteration related to skarn deposits is the occurrence of hairline fractures and stylolites in sedimentary rocks beyond the limit of calc-silicate minerals. This has been documented for many different skarn deposits (e.g. Meinert et al., 1997) but is particularly common in Zn skarns.
Most molybdenum skarns are associated with leucocratic granites and range from high grade, relatively small deposits (Azegour, Morocco, Permingeat, 1957) to low grade, bulk tonnage deposits (Little Boulder Creek, Idaho, Cavanaugh, 1978). Numerous small occurrences are also found in Precambrian stable cratons associated with pegmatite, aplite, and other leucocratic rocks (Vokes, 1963). Most molybdenum skarns contain a variety of metals including W, Cu, Zn, Pb, Bi, Sn, and U and some are truly polymetallic in that several metals need to be recovered together in order for the deposits to be mined economically. Mo-W-Cu is the most common association and some tungsten skarns and copper skarns contain zones of recoverable molybdenum.
Most molybdenum skarns occur in silty carbonate or calcareous clastic rocks; Cannivan Gulch, Montana (Darling, 1990) is a notable exception in that it occurs in dolomite. Hedenbergitic pyroxene is the most common calc-silicate mineral reported from molybdenum skarns with lesser grandite garnet (with minor pyralspite component), wollastonite, amphibole, and fluorite. This skarn mineralogy indicates a reducing environment with high fluorine activities. These deposits have not received significant study outside of the Soviet Union and there has not been a modern review since the brief summary by Einaudi et al. (1981).
Tin skarns are almost exclusively associated with high-silica granites generated by partial melting of continental crust. Major reviews of tin skarn deposits include Einaudi et al. (1981) and Kwak (1987). Tin skarns can be subdivided according to several criteria including proximal versus distal, calcic versus magnesian, skarn-rich versus skarn-poor, oxide-rich versus sulfide-rich, and greisen versus skarn. Unfortunately, few of these categories are mutually exclusive.
Many large tin skarn systems are zoned spatially from skarn-rich to skarn-poor (or absent). For example, in the Renison Bell area of Tasmania, Australia there is a single large magmatic/hydrothermal system zoned from a proximal calcic tin skarn with minor cassiterite disseminated in a sulfide-poor garnet-pyroxene gangue to a distal magnesian massive sulfide replacement body containing abundant cassiterite and a complete absence of calc-silicate minerals. The distal massive sulfide ore body (Renison Bell) is a major ore deposit and the proximal skarn body (Pine Hill) has not and probably never will be mined.
Einaudi et al. (1981) emphasized that there is a common thread linking the several types of tin skarn deposits and that is the characteristic suite of trace elements (Sn, F, B, Be, Li, W, Mo, and Rb) in the ore and in associated igneous rocks. This suite distinguishes tin skarns from all other skarn types. Kwak (1987) makes a further distinction in that many tin skarn deposits develop a greisen alteration stage which is superimposed upon the intrusion, early skarn, and unaltered carbonate. Greisen alteration is characterized by high fluorine activities and the presence of minerals like fluorite, topaz, tourmaline, muscovite, grunerite, ilmenite, and abundant quartz. In many cases this greisen-stage alteration completely destroys earlier alteration stages. Of particular importance, greisen-style alteration is absent from all other skarn types.
There are several mineralogical features of tin skarns that should be highlighted. From a mining standpoint, the most important is that tin can be incorporated into silicate minerals, such as garnet, sphene, and idocrase, where it is economically unrecoverable. Dobson (1982) reports garnet containing up to 6% Sn in skarn at Lost River, Alaska. Thus, large deposits such as Moina in Tasmania (Kwak and Askins, 1981), can contain substantial amounts of tin that cannot be recovered with present or foreseeable technology. Extensive retrograde or greisen alteration of early tin-bearing skarn minerals can liberate this tin and cause it to precipitate in oxide or sulfide ore. Thus, the skarn destructive stages of alteration are particularly important in tin skarn deposits. As noted by Kwak (1987), the most attractive ore bodies occur in the distal portions of large skarn districts where massive sulfide or oxide replacements occur without significant loss of tin in calc-silicate minerals like garnet.
There are many other types of skarn which historically have been mined or explored for a variety of metals and industrial minerals. Some of the more interesting include rare metal and rare earth element enriched skarns (e.g. Kato, 1989; Birkett and SInclair, 1998). REEs tend to be enriched in specific mineral phases such as garnet, idocrase, epidote, and allanite. Vesuvianite and epidote with up to 20% REE (Ce>La>Pr>Nd) have been found in some gold skarns and zinc skarns (Gemmel et al., 1992; Meinert, unpublished data).
Some skarns contain economic concentrations of REEs and uranium (Kwak and Abeysinghe, 1987; Lentz, 1991, 1998). The Mary Kathleen skarn deposit in Queensland, Australia is unusual in that REEs and uranium daughter minerals in fluid inclusions suggest that these elements can be strongly concentrated in high-temperature hydrothermal fluids (Kwak and Abeysinghe, 1987). This suggests that other metasomatic environments should be examined for possible concentrations of REEs and uranium.
The occurrence of platinum group elements is reported in some skarns (e.g. Knopf, 1942; Korobeynikov et al., 1998). These deposits have not been well documented in the literature and most appear to represent metasomatism of ultramafic rocks (e.g. Yu, 1985). It is difficult to evaluate the abundance of PGEs in different skarn types because PGEs have not been routinely analyzed until recently. Geochemical considerations suggest that PGEs could be transported under very acidic, oxidized conditions (Wood, 1989). In the skarn environment such conditions might be reached in the greisen alteration stage of tin skarns. This might be a direction for future research and exploration.
In most skarns there is a general zonation pattern of proximal garnet, distal pyroxene, and idocrase (or a pyroxenoid such as wollastonite, bustamite, or rhodonite) at the contact between skarn and marble. In addition, individual skarn minerals may display systematic color or compositional variations within the larger zonation pattern. For example, proximal garnet is commonly dark red-brown, becoming lighter brown and finally pale green near the marble front (e.g. Atkinson and Einaudi, 1978). The change in pyroxene color is less pronounced but typically reflects a progressive increase in iron and/or manganese towards the marble front (e.g. Harris and Einaudi, 1982). For some skarn systems, these zonation patterns can be "stretched out" over a distance of several kilometres and can provide a significant exploration guide (e.g. Meinert, 1987). Details of skarn mineralogy and zonation can be used to construct deposit-specific exploration models as well as more general models useful in developing grass roots exploration programs or regional syntheses. Reasonably detailed zonation models are available for copper, gold, and zinc skarns (Meinert, 1997). Other models can be constructed from individual deposits which have been well studied such as the Hedley Au skarn (Ettlinger, 1992; Ray et al., 1993) or the Groundhog Zn skarn (Meinert, 1982).
Skarn formation spans almost the complete range of potential ore-forming environments. Most geochemical studies of skarn deposits have focused on mineral phase equilibria, fluid inclusions, isotopic investigations of fluid sources and pathways, and determination of exploration anomaly and background levels. Experimental phase equilibria studies are essential for understanding individual mineral reactions. Such studies can be extended using thermodynamic data to include variable compositions). Another approach is to use a self-consistent thermodynamic database to model potential skarn-forming solutions (e.g. Flowers and Helgeson, 1983; Johnson and Norton, 1985; Ferry and Baumgartner, 1987). Fractionation of elements between minerals (e.g. Ca:Mg in carbonate, Bowman et al., 1982; Bowman and Essene, 1984) also can be used to estimate conditions of skarn formation. A general review of phase equilibria applicable to skarn systems is presented by Bowman (1998). A more specialized treatment of the vector representation of skarn mineral stabilities is presented by Burt (1998). Recent work has incorporated standard phase equilbria treatment of skarn mineralogy along with fluid dynamics to model the metasomatic evolution of skarn systems (Dipple and Gerdes, 1998).
Fluid inclusion studies of many ore deposit types focus on minerals such as quartz, carbonate, and fluorite which contain numerous fluid inclusions, are relatively transparent, and are stable over a broad T-P-X range. However, this broad T-P-X range can cause problems in interpretation of fluid inclusion data, because these minerals may grow and continue to trap fluids from early high temperature events through late low temperature events (Roedder, 1984). In contrast, high temperature skarn minerals such as forsterite, diopside, etc. are unlikely to trap later low temperature fluids (beyond the host mineral's stability range) without visible evidence of alteration. Thus, fluid inclusions in skarn minerals provide a relatively unambiguous opportunity to measure temperature, pressure, and composition of skarn-forming fluids.
Much of the skarn fluid inclusion literature prior to the mid-1980's has been summarized by Kwak (1986), especially studies of Sn and W skarn deposits. Such studies have been very useful in documenting the high temperatures (>700¡C) and high salinities (>50 wt. % NaCl equiv. and multiple daughter minerals) which occur in many skarns. All the skarn types summarized in Meinert (1992) have fluid inclusion homogenization temperatures up to and exceeding 700¡C except for copper and zinc skarns, deposits in which most fluid inclusions are in the 300-550¡C range. This is consistent with the relatively shallow and distal geologic settings inferred respectively for these two skarn types.
Salinities in most skarn fluid inclusions are high; documented daughter minerals in skarn minerals include NaCl, KCl, CaCl2, FeCl2, CaCO3, CaF2, C, NaAlCO3(OH)2, Fe2O3, Fe3O4, AsFeS, CuFeS2, and ZnS (Table 2). Haynes and Kesler (1988) describe systematic variations in NaCl:KCl:CaCl2 ratios in fluid inclusions from different skarns reflecting differences in the fluid source and the degree of mixing of magmatic, connate, and meteoric fluids. In general, magmatic fluids have KCl>CaCl2 whereas high-CaCl2 fluids appear to have interacted more with sedimentary wall rocks.
Fluid inclusions can provide direct evidence for the content of CO2 (both liquid and gas), CH4, N2, H2S and other gases in hydrothermal fluids. Studies of gas phases and immiscible liquids in fluid inclusions typically show a dominance of CO2, a critical variable in skarn mineral stability. Although no comparative studies have been done, it appears that CH4 is slightly more abundant than CO2 in reduced systems like tungsten skarns (Fonteilles et al., 1989; Gerstner et al., 1989) whereas CO2 is more abundant than CH4 in more oxidized systems like copper and zinc skarns (Megaw et al., 1988).
Studies of fluid inclusions in specific skarn mineral phases are particularly useful in documenting the temporal and spatial evolution of skarn-forming fluids and how those changes correlate with compositional, experimental, and thermodynamic data (e.g. Kwak and Tan, 1981; Meinert, 1987). Fluid inclusions also provide direct evidence for the temperature and salinity shift in most skarn systems between prograde and retrograde skarn events. For example, most garnet and pyroxene fluid inclusions in iron skarns have homogenization temperatures of 370->700¡C and 300-690¡C, respectively, with salinities up to 50 wt. % NaCl equivalent, whereas retrograde epidote and crosscutting quartz veins have homogenization temperatures of 245-250¡C and 100-250¡C, respectively, with salinities of less than 25 wt. % NaCl equivalent.
In gold skarns, prograde garnet and pyroxene homogenization temperatures are up to 730¡C and 695¡C, respectively, with salinities up to 33 wt. % NaCl equivalent. In contrast, scapolite, epidote, and actinolite from these skarns have homogenization temperatures of 320-400¡C, 255-320¡C, and 320-350¡C, respectively. In tungsten skarns, prograde garnet and pyroxene homogenization temperatures are up to 800¡C and 600¡C, respectively, with salinities up to 52 wt. % NaCl equivalent. In contrast, amphibole and quartz from these skarns have homogenization temperatures of 250-380¡C and 290-380¡C, respectively with salinities of 12-28 and 2.5-10.5 wt. % NaCl equivalent (data summarized in Meinert, 1992).
Isotopic investigations, particularly the stable isotopes of C, O, H, and S, have been critically important in documenting the multiple fluids present in most large skarn systems (Shimazaki, 1988; Bowman, 1998). The pioneering study of Taylor and O'Neill (1977) demonstrated the importance of both magmatic and meteoric waters in the evolution of the Osgood Mountain W skarns. Bowman et al. (1985) demonstrated that in high temperature W skarns, even some of the hydrous minerals such as biotite and amphibole can form at relatively high temperatures from water with a significant magmatic component (see also Marcke de Lummen, 1988).
Specifically, garnet, pyroxene, and associated quartz from the skarn deposits summarized in Meinert (1992) all have ¶18O values in the +4 to +9 range consistent with derivation from magmatic waters. In contrast, ¶18O values for sedimentary calcite, quartz, and meteoric waters in these deposits are distinctly different. In most cases, there is a continuous mixing line between original sedimentary ¶18O values and calculated ¶18O values for magmatic hydrothermal fluids at the temperatures of prograde skarn formation. Similar mixing is indicated by ¶13C values in calcite, ranging from typical sedimentary ¶13C values in limestone away from skarn to typical magmatic values in calcite interstitial to prograde garnet and pyroxene (Brown et al., 1985). Hydrous minerals such as biotite, amphibole, and epidote from different skarn deposits also display ¶18O and ¶D values ranging from magmatic to local sedimentary rocks and meteoric waters (Layne et al., 1991). Again, mixing of multiple fluid sources is indicated.
Sulfur isotopic studies on a variety of sulfide minerals (including pyrite, pyrrhotite, molybdenite, chalcopyrite, sphalerite, bornite, arsenopyrite, and galena) from the skarn deposits summarized in Table 2 indicate a very narrow range of ¶34 values, consistent with precipitation from magmatic fluids. For some of the more distal zinc skarns, sulfur isotopic studies indicate that the mineralizing fluids acquired some of their sulfur from sedimentary rocks (including evaporites) along the fluid flow path (Megaw et al., 1988).
Overall, stable isotopic investigations are consistent with fluid inclusion and mineral equilibria studies which demonstrate that most large skarn deposits form from diverse fluids, including early, high temperature, highly saline brines directly related to crystallizing magma systems (e.g. Auwera and Andre, 1988). In many systems, the highest salinity fluids are coincident with peak sulfide deposition. In addition, at least partial mixing with exchanged connate or meteoric fluids is required for most deposits with the latest alteration events forming largely from dilute meteoric waters.
Even though skarn metal contents are quite variable, anomalous concentrations of pathfinder elements in distal skarn zones can be an important exploration guide. Geochemical studies of individual deposits have shown that metal dispersion halos can be zoned from proximal base metal assemblages, through distal precious metal zones, to fringe Pb-Zn-Ag vein concentrations (e.g. Theodore and Blake, 1975). Anomalies of 10s to 100s of ppm for individual metals can extend for more than 1000 meters beyond proximal skarn zones. Comparison of geochemical signatures among different skarn classes suggests that each has a characteristic suite of anomalous elements and that background levels for a particular element in one skarn type may be highly anomalous in other skarns. For example, Au, Te, Bi, and As values of 1, 10, 100, and 500 ppm, respectively, are not unusual for gold skarns but are rare to absent for other skarn types (e.g. Meinert et al., 1990; Myers and Meinert, 1991).
Some skarns have a strong geophysical response (Chapman and Thompson, 1984; Emerson, 1986). Almost all skarns are significantly denser than the surrounding rock and therefore may form a gravitational anomaly or seismic discontinuity. This is particularly evident in some of the large iron skarns which may contain more than a billion tons of magnetite (specific gravity, 5.18). In addition, both skarns and associated plutons may form magnetic anomalies (Spector, 1972). Relatively oxidized plutons typically contain enough primary magnetite to form a magnetic high whereas reduced plutons typically contain ilmenite rather than magnetite and may form a magnetic low (Ishihara, 1977). Skarns may form a magnetic high due to large concentrations of magnetite (Chapman et al., 1986) or other magnetic minerals such as high temperature pyrrhotite (Wotruba et al., 1988). Since metasomatism of dolomitic rocks tends to form abundant magnetite, in magnesian skarn deposits a strong magnetic signature may be able to distinguish original protolith as well as the presence of skarn (Hallof and Winniski, 1971; Chermeninov, 1988).
Electrical surveys of skarns need to be interpreted carefully. Either disseminated or massive sulfide minerals may give strong IP, EM, or magnetotelluric responses in skarn (Emerson and Welsh, 1988). However, metasomatism of carbonate rock necessarily involves the redistribution of carbon. The presence of carbonaceous matter, especially if in the form of graphite, can strongly effect electrical surveys. Such carbon-induced anomalies may be distant from or unrelated to skarn ore bodies.
A few skarns contain sufficient uranium and thorium to be detectable by airborne or ground radiometric surveys (e.g. Mary Kathleen, Australia, Kwak and Abeysinghe, 1987). Detailed studies of such deposits demonstrate that relatively small skarns can be detected and that different types of skarns can be distinguished (e.g. Lentz, 1991). Although gravity, magnetic, electrical, and radiometric methods have all been applied to skarn deposits, their use has not been widespread. Because of the variability of skarn deposits, it probably is necessary to tailor specific geophysical methods to individual skarn deposits or types.
Most major skarn deposits are directly related to igneous activity and broad correlations between igneous composition and skarn type have been described by several workers (Zharikov, 1970; Shimazaki, 1975,1980; Einaudi et al., 1981; Kwak and White, 1982; Meinert, 1983; Newberry and Swanson, 1986; Newberry, 1987; 1990). Averages of large amounts of data for each skarn type can be summarized on a variety of compositional diagrams to show distinctions among skarn classes. Tin and molydenum skarns typically are associated with high silica, strongly differentiated plutons. At the other end of the spectrum, iron skarns usually are associated with low silica, iron-rich, relatively primitive plutons. Such diagrams are less useful for detailed studies, however, because of the wide range of igneous compositions possible for an individual skarn deposit and the difficulty of isolating the effects of metasomatism and late alteration.
Other important characteristics include the oxidation state, size, texture, depth of emplacement, and tectonic setting of individual plutons. For example, tin skarns are almost exclusively associated with reduced, ilmenite-series plutons which can be characterized as S-type or anorogenic. These plutons tend to occur in stable cratons in which partial melting of crustal material may be instigated by incipient rifting. Many gold skarns are also associated with reduced, ilmenite-series plutons. However, gold skarn plutons typically are mafic, low-silica bodies which could not have formed by melting of sedimentary crustal material. In contrast, plutons associated with copper skarns, particularly porphyry copper deposits, are strongly oxidized, magnetite-bearing, I-type and associated with subduction-related magmatic arcs. These plutons tend to be porphyritic and emplaced at shallow levels in the earthÕs crust. Tungsten skarns, on the other hand, are associated with relatively large, coarse-grained, equigranular plutons or batholithic complexes indicative of a deeper environment.
Tectonic setting, petrogenesis, and skarn deposits are intimately intertwined. Some modern textbooks use tectonic setting to classify igneous provinces (Wilson, 1989) or different kinds of ore deposits (Sawkins, 1984). This approach has been less successful in describing ore deposits such as skarns which are the result of processes that can occur in almost any tectonic setting. A useful tectonic classification of skarn deposits should group skarn types which commonly occur together and distinguish those which typically occur in specialized tectonic settings. For example, calcic Fe-Cu skarn deposits are virtually the only skarn type found in oceanic island-arc terranes. Many of these skarns are also enriched in Co, Ni, Cr, and Au. In addition, some economic gold skarns appear to have formed in back arc basins associated with oceanic volcanic arcs (Ray et al., 1988). Some of the key features that set these skarns apart from those associated with more evolved magmas and crust are their association with gabbroic and dioritic plutons, abundant endoskarn, widespread sodium metasomatism, and the absence of Sn and Pb. Collectively, these features reflect the primitive, oceanic nature of the crust, wall rocks, and plutons.
The vast majority of skarn deposits are associated with magmatic arcs related to subduction beneath continental crust. Plutons range in composition from diorite to granite although differences among the main base metal skarn types appear to reflect the local geologic environment (depth of formation, structural and fluid pathways) more than fundamental differences of petrogenesis (Nakano et al., 1990). In contrast, gold skarns in this environment are associated with particularly reduced plutons that may represent a restricted petrologic history.
The transition from subduction beneath stable continental crust to post-subduction tectonics is not well understood. Magmatism associated with shallow subduction angles may have more crustal interaction (Takahashi et al., 1980) and floundering of the downgoing slab may result in local rifting. During this stage the magmatic arc may widen or migrate further inland. Plutons are granitic in composition and associated skarns are rich in Mo or W-Mo with lesser Zn, Bi, Cu, and F. Many of these skarns are best described as polymetallic with locally important Au and As.
Some skarns are not associated with subduction-related magmatism. These skarns may be associated with S-type magmatism following a major period of subduction or they may be associated with rifting of previously stable cratons. Plutons are granitic in composition and commonly contain primary muscovite and biotite, dark gray quartz megacrysts, miarolitic cavities, greisen-type alteration, and anomalous radioactivity. Associated skarns are rich in tin or fluorine although a host of other elements are usually present and may be of economic importance. This evolved suite includes W, Be, B, Li, Bi, Zn, Pb, U, F, and REE.
This section currently is in the form of an oral presentation given during the 1996-1997 SEG Thayer Lindsley lecture tour and during short courses for mining/exploration companies. When time becomes available it will be adapted for presentation here.
Comments and questions concerning this page can be directed to:
Larry Meinert (email: Lmeinert@email.smith.edu)
Department of Geology
Northampton, MA 01063
November 2, 2007
© 1995 (Larry Meinert)