Ore Geology Reviews 40 (2011) 1–26

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Ore Geology Reviews

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Review

Magmatic to hydrothermal metal fluxes in convergent and collided margins

Jeremy P. Richards ⁎Dept. of Earth and Atmospheric Sciences, University of Alberta, Edmonton, Alberta, Canada T6G 2E3

⁎ Tel.: +1 780 492 3430; fax: +1 780 492 2030.E-mail address: Jeremy.Richards@ualberta.ca.

0169-1368/$ – see front matter © 2011 Elsevier B.V. Aldoi:10.1016/j.oregeorev.2011.05.006

a b s t r a c t

a r t i c l e i n f o

Article history:Received 1 December 2010Received in revised form 18 May 2011Accepted 19 May 2011Available online 27 May 2011

Keywords:Porphyry depositEpithermal depositSubductionPost-subductionMagmatic–hydrothermal fluidOre formation

Metals such as Cu, Mo, Au, Sn, and W in porphyry and related epithermal mineral deposits are derivedpredominantly from the associatedmagmas, viamagmatic–hydrothermal fluids exsolved upon emplacement intothemid- to upper crust. Fourmain sources exist formagmas, and thereforemetals, in convergent and collided platemargins: the subducting oceanic plate basaltic crust, subducted seafloor sediments, the asthenospheric mantlewedge between the subducting and overriding plates, and the upper plate lithosphere. This paper firstly examinesthe source of normal arc magmas, and concludes that they are predominantly derived from partial melting of themetasomatized mantle wedge, with possible minor contributions from subducted sediments. Although somemetals may be transferred from the subducting slab via dehydration fluids, the bulk of the metals in the resultantmagmas are probably derived from the asthenospheric mantle. The most important contributions from the slabfromametallogenicperspective areH2O, S, andCl, aswell as oxidants. Partialmeltingof the subductedoceanic crustand/or sediments may occur under some restricted conditions, but is unlikely to be a widespread process (inPhanerozoic arcs), and does not significantly differ metallogenically from slab-dehydration processes.Primary, mantle-derived arc magmas are basaltic, but differ frommid-ocean ridge basalt in having higher watercontents (~10× higher), oxidation states (~2 log fO2 units higher), and concentrations of incompatible elementsand other volatiles (e.g., S and Cl). Concentrations of chalcophile and siderophile metals in these partial meltsdepend critically on the presence and abundance of residual sulfide phases in the mantle source. At relativelyhigh abundances of sulfides thought to be typical of active arcs where fS2 and fO2 are high (magma/sulfideratio=102–105), sparse, highly siderophile elements such as Au and PGE will be retained in the source, butmagmas may be relatively undepleted in abundant, moderately chalcophile elements such as Cu (and perhapsMo). Suchmagmas have the potential to form porphyry Cu±Modeposits upon emplacement in the upper crust.Gold-rich porphyry deposits would only form where residual sulfide abundance was very low (magma/sulfideratio N105), perhaps due to unusually high mantle wedge oxidation states.In contrast, some porphyry Mo and all porphyry Sn–W deposits are associated with felsic granitoids, derivedprimarily from melting of continental crust during intra-plate rifting events. Nevertheless, mantle-derivedmagmas may have a role to play as a heat source for anatexis and possibly as a source of volatiles and metals.In post-subduction tectonic settings Tulloch and Kimbrough, 2003, such as subduction reversal or migration, arccollision, continent–continent collision, and post-collisional rifting, a subducting slab source no longer exists,andmagmas are predominantly derived from partialmelting of the upper plate lithosphere. This lithospherewillhave undergone significant modification during the previous subduction cycle, most importantly with theintroduction of large volumes of hydrous, mafic (amphibolitic) cumulates residual from lower crustaldifferentiation of arc basalts. Small amounts of chalcophile and siderophile element-rich sulfidesmay also be leftin these cumulates. Partial melting of these subduction-modified sources due to post-subduction thermalreadjustments or asthenospheric melt invasion will generate small volumes of calc-alkaline to mildly alkalinemagmas, which may redissolve residual sulfides. Such magmas have the potential to form Au-rich as well asnormal Cu±Moporphyry and epithermal Au systems, depending on the amounts of sulfide present in the lowercrustal source. Alkalic-type epithermal Au deposits are an extreme end-member of this range of post-subductiondeposits, formed from subduction-modified mantle sources in extensional or transtensional environments.Ore formation in porphyry and related epithermal environments is critically dependent on the partitioning ofmetals from the magma into an exsolving magmatic–hydrothermal fluid phase. This process occurs mostefficiently at depths greater than ~6 km, within large mid- to upper crustal batholithic complexes fed by arc orpost-subduction magmas. Under such conditions, metals will partition efficiently into a single-phase,supercritical aqueous fluid (~2–13 wt.% NaCl equivalent), which may exist as a separate volatile plume or asbubbles entrained in buoyant magma. Focusing of upward flow of bubbly magma and/or fluid into the apicalregions of the batholithic complex forms cupolas, which represent high mass- and heat-flux channelways

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2 J.P. Richards / Ore Geology Reviews 40 (2011) 1–26

towards the surface. Cupolas may be self-organizing to the extent that once formed, further magma and fluidflow will be enhanced along the weakened and heated axes. Cupolas may form initially as breccia pipes byvolatile phase (rather than magma) reaming-out of extensional structures in the brittle cover rocks, to befollowed immediately by magma injection to form cylindrical plugs or dikes.Cupola zonesmay extend to surface,wheremagmas andfluids vent as volcanic products and fumaroles. Betweenthe surface and the underlyingmagma chamber, a very steep thermal gradient exists (700°–800 °C over b5 km),which is the primary cause of vertical focusing of oremineral deposition. The bulk ofmetals (Cu±Mo±Au) thatformsporphyry ore bodies areprecipitatedover anarrowtemperature interval between~425° and320 °C,whereisotherms in the cupola zone rise to within ~2 km of the surface. Over this temperature range, four importantphysical and physicochemical factors act to maximize ore mineral deposition: (1) silicate rocks transition fromductile to brittle behavior, thereby greatly enhancing fracture permeability and enabling a threefold pressuredrop; (2) silica shows retrograde solubility, thereby further enhancing permeability and porosity for oredeposition; (3) Cu solubility dramatically decreases; and (4) SO2 dissolved in the magmatic–hydrothermal fluidphase disproportionates to H2S and H2SO4, leading to sulfide and sulfate mineral deposition and the onset ofincreasingly acidic alteration.The bulk of the metal flux into the porphyry environment may be carried by moderately saline supercriticalfluids or vapors, with a volumetrically lesser amount by saline liquid condensates. However, these vapors rapidlybecome dilute at lower temperatures and pressures, such that they lose their capacity to transport metals aschloride complexes. They retain significant concentrations of sulfur species, however, and bisulfide complexingof Cu and Au may enable their continued transport into the epithermal regime. In the high-sulfidationepithermal environment, intense acidic (advanced-argillic) alteration is caused by the flux of highly acidicmagmatic volatiles (H2SO4, HCl) in this vapor phase. Ore formation, however, is paragenetically late, andmay belocated in these extremely altered and leached cap rocks largely because of their high permeability and porosity,rather than there being any direct genetic connection. Ore-forming fluids, where observed, are low- tomoderate-salinity liquids, and are thought to represent later-stage magmatic–hydrothermal fluids that haveascended along shallower (cooler) geothermal gradients that either do not, or barely, intersect the liquid–vaporsolvus. Such fluids “contract” from the original supercritical fluid or vapor to the liquid phase. Brief intersectionof the liquid–vapor solvus may be important in shedding excess chloride and chloride-complexed metals (suchas Fe), so that bisulfide-complexed metals remain in solution. Such a restrictive pressure–temperature path islikely to occur only transiently during the evolution of amagmatic–hydrothermal system,whichmay explain therarity of high-sulfidation Cu–Au ore deposits, despite the ubiquitous occurrence of advanced-argillic alteration inthe lithocaps above porphyry-type systems.

© 2011 Elsevier B.V. All rights reserved.

Contents

1. Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 32. Magma generation in convergent and collided margins: geochemical characteristics and partitioning of metals . . . . . . . . . . . . . . . . . . 3

2.1. Slab dehydration and asthenospheric melting in subduction zones . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 32.1.1. Behavior of metals. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4

2.2. Sediment dehydration and/or melting in subduction zones . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 52.2.1. Behavior of metals. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5

2.3. Oceanic slab melting in subduction zones . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 62.3.1. Behavior of metals. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6

2.4. Supra-subduction zone lithospheric melting: the MASH process . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 62.4.1. Behavior of metals. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 82.4.2. Sources of Mo . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8

2.5. Lithospheric melting during post-subduction events . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 82.5.1. Behavior of metals in subduction-modified sources . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 9

2.6. Crustal melting during post-collisional stress relaxation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 102.6.1. Sources of metals . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 10

3. Behavior of metals during magma fractionation and fluid exsolution in the upper crust . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 113.1. Partitioning of metals from magma into exsolving hydrothermal fluid. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 11

4. Magmatic–hydrothermal ore formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 134.1. Porphyry Cu ore formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 134.2. Epithermal Cu–Au ore formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 15

4.2.1. High-sulfidation epithermal Cu–Au deposits. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 154.2.2. Low-sulfidation epithermal Au deposits (including alkalic-type deposits) . . . . . . . . . . . . . . . . . . . . . . . . . . . . 17

5. Summary and conclusions. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 185.1. Sources of magmas and metals . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 185.2. Porphyry and epithermal ore formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 18

Acknowledgments . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 19References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 19

3J.P. Richards / Ore Geology Reviews 40 (2011) 1–26

1. Introduction

The question of the source of various elements in convergent andcollidedmarginmagmas has challenged geologists for decades. Igneouspetrologists seek to understand the petrogenesis of such magmasthrough geochemical and isotopic tracing,whereas economic geologistsare generally more interested in the source of potentially valuableelements suchasCu,Mo, Sn,W,Au, andplatinumgroupelements (PGE),which may ultimately be found in intrusion-related hydrothermaldeposits.

Igneous petrologists are broadly in agreement that arc magmas areprimarily derived from hydrous melting of the asthenospheric mantlewedge above subducting plates, but melts from the subducted oceaniccrust (including sediments) and the upper plate lithosphere may alsobe involved to varying degrees.

Economic geologists are also broadly in agreement that ore-formingelements are partitioned from such magmas into an exsolving volatilephase upon emplacement in the upper crust, and may then beprecipitated from thesefluids during cooling, fluidmixing, andwallrockreaction processes in porphyry-type and related epithermal mineraldeposits. However, these process theories do not address where themetals originally came from, nor why porphyry deposits vary sowidelyin their metal contents (from Au-rich, through Cu±Mo±Au, to Mo-only deposits, with Sn–W deposits forming a distinct variant).

In addition to subduction-related calc-alkaline magmas, a diversesuite of calc-alkaline to alkaline magmas is generated in post-subduction and collisional tectonic settings, and these magmaticsystems may also generate porphyry and epithermal ore deposits.Such systems raise an additional set of petrogenetic and metallogenicquestions.

It is the intent of this paper to merge these different geologicalperspectives on magmagenesis and metallogeny in order to discussprimary metal fluxes in convergent and collisional margins in terms ofigneous petrogenetic and magmatic–hydrothermal processes. Theultimate metal inventory and metal ratios in any given porphyry orrelated deposit is secondarily controlled by late-stage magmatic andshallow crustal processes. These processes are examined, closingwith areview of fluid and metal sources and behavior in related epithermalenvironments.

2. Magma generation in convergent and collided margins:geochemical characteristics and partitioning of metals

Most magmas erupted through or emplaced within the Earth's crustare not primary magmas (in the sense of being chemically unmodifiedsince extraction from their source), andmost are not even primitive (inthe sense of being relatively unevolved; Hildreth and Moorbath, 1988;Leeman, 1983; Neuendorf et al., 2005; Smith et al., 2010; Thirlwall et al.,1996). Except formagmasproduced and erupted in extensional tectonicregimes (where rapid ascent to the surface is facilitated by normalfaulting), most deeply-derived magmas undergo some degree offractionation and crustal contamination during their passage towardsthe surface. It is therefore challenging to isolate geochemical andisotopic characteristics ofmagma source regions from the effects of laterprocesses (Davidson, 1996). Magmas erupted through mature conti-nental crust are the most difficult to fingerprint uniquely in terms ofsource characteristics because wallrock assimilation and fractionalcrystallization (AFC; DePaolo, 1981) are ubiquitous and commonlyextensiveprocesses thatwill significantlymodify bulk rock geochemicaland isotopic compositions; and yet, these are also the magmas that aremost commonly associated with porphyry- and epithermal-typemineral deposits. The difficulty in constraining source characteristicsin such magmas is perhaps responsible for the plethora of theories thathave beenproposed for the origin of ore-formingmagmas in convergentmargin settings, ranging from the melting of subducting oceanic crustand/or seafloor sediments, through melting of subduction metasoma-

tized asthenospheric or lithospheric mantle, to melting of underplatedor primitive lower crustal rocks, and even melting of evolved crustalrocks in the case of some felsic porphyry Mo and Sn–W magmas.

Therefore, rather than start by trying to identify a unique sourcefor the typical intermediate-to-felsic calc-alkaline magmas that areassociated with ore deposits in mature convergent margins, I beginthis review by focusing on the much better constrained primitiveisland arc environment, where the effects of fractionation and crustalcontamination, particularly by continentally derived materials, areminimized, and processes in mantle source regions can be moreclearly defined.

2.1. Slab dehydration and asthenospheric melting in subduction zones

There is a general consensus that, with the exception of youngoceanic lithosphere (b25 m.y.-old; Defant and Drummond, 1990) orplate edges (Yogodzinski et al., 2001), basaltic oceanic crust undergoeslow-temperature, high pressure metamorphism upon subduction,which releasesfluids through a series of prograde dehydration reactionsto form anhydrous eclogite (Fig. 1). Water and other volatilecomponents and solutes (including S and Cl) were originally incorpo-rated into oceanic crustal and upper mantle rocks during oxidizingseafloor alteration, generatinghydrousminerals suchas serpentine, talc,amphibole, micas, chlorite, zoisite, chloritoid, and lawsonite. Variousexperimental studies have shown that these minerals undergodehydration reactions over a depth range extending to ~100 km,corresponding to the blueschist–eclogite transition in crustal rocks;serpentine (antigorite) and the 10 Å equivalent of chlorite may extendthis range to200 km(e.g., Dvir et al., 2011; Forneris andHolloway, 2003;Fumagalli and Poli, 2005; Poli and Schmidt, 2002; Schmidt and Poli,1998; Ulmer and Trommsdorff, 1995). Below these depths, theanhydrous eclogitic crust is essentially infusible, and the dense slabcontinues its descent into the mantle without melting. See reviews ofthis subject by Richards (2003, 2005, and references therein).

Numerous studies have explored the character of the fluids that arereleased from the dehydrating slab, because they are thought to accountfor the unique geochemical character of subduction-related magmasduring later partial melting in the metasomatized asthenosphericmantle wedge (located between the downgoing slab and the upperplate). Slab-derived fluids are thought to be water-rich at these depths,and to carry significant amounts of other volatile components such as Cland S. For example, salinities in the range of 4–10 wt.% NaCl have beeninferred from primitive basalt or melt inclusion studies (de Hoog et al.,2001; Kent et al., 2002; Portnyagin et al., 2007; Wallace, 2005;Wysoczanski et al., 2006), and salinities of 0.4–2 wt.% NaCl equivalentwere measured in fluid inclusions from high-pressure rocks thought torepresent subductedoceanicmantle (Scambelluri et al., 2004). At higherpressures (~6 GPa) and greater depths (~175 km) there may no longerbe a physical distinction between solute-rich aqueous liquids andhydrous silicate melts, and the fluid may be supercritical in nature (e.g.,Kessel et al., 2005a,b), but the role of such deeply released fluids insubduction zone magmatism is unclear (see discussion in Richards andKerrich, 2007).

In addition to volatiles, water-soluble large ion lithophile elements(LILE: K, Rb, Cs, Ca, Sr, andBa, andU), andB, Pb, As, andSb (Breeding et al.,2004; Hattori et al., 2005; Hattori and Guillot, 2003; Kogiso et al., 1997;Manning, 2004; Tatsumi et al., 1986) are thought to bemobilized into theforearc mantle wedge by dehydration fluids, which may then beconvected by corner-flow into the sub-arc melting zone (Fig. 1). Thesefluid–mobile components are also characteristically enriched in arcmagmas (e.g., typical ranges of: 1–3 wt.% H2O, 500–2000 ppm Cl, 900–2500 ppm S; Davidson, 1996; Gill, 1981; Noll et al., 1996; Sobolev andChaussidon, 1996;Wallace, 2005; Portnyagin et al., 2007),which is takenas evidence of aqueousfluidmetasomatism of themantlewedgemagmasource. Silica may also be significantly mobilized in these slab fluids(Aerts et al., 2010;Manning, 2004), aswell as Tl and Cu (Noll et al., 1996;

Mantle corner flow

Sea level

Asthenosphere

Volcanic arc

Talc

Serpentine

Chlorite +serpentine

Chlorite

1000°C

600°C

Oce

an

ic m

an

tle

lith

osp

he

reOceanic crust

Dehydration of oceanic

lithosphere1000°C

Amphibole

ZoisiteCtd

Oceanic crust

Partial melting

Mantlelithosphere

Chlorite–10Å phase

+ serpentine

0 km

100 km

200 km

1400°C

1400°C

600°C

1000°C

Eclogite

Sediment

Metasomatizedasthenosphere

Fig. 1. Structure and processes beneath an oceanic island arc (sources: Tatsumi and Eggins, 1995; Schmidt and Poli, 1998; Winter, 2001; Poli and Schmidt, 2002; Fumagalli and Poli,2005). Primary hydrous basaltic arc magmas are derived from partial melting of the metasomatized asthenospheric mantle wedge. Mineral zones shown in the subducting plateindicate lower limits of stability of hydrous phases in the basaltic oceanic crust and peridotitic mantle lithosphere. Abbreviation: Ctd=chloritoid.

4 J.P. Richards / Ore Geology Reviews 40 (2011) 1–26

Stolper and Newman, 1994). The normally relatively incompatible highfield strength elements (HFSE) Ti, Nb, and Ta are not significantly fluidsoluble under subduction zone conditions, and are retained in mineralssuch as rutile either in the slab or the mantle wedge (Audétat andKeppler, 2005; Brenan et al., 1994; Green and Adam, 2003). Arc magmasderived from these sources therefore show characteristic negativeanomalies for these three elements on mantle-normalized spiderdiagrams, but display enrichments in most other incompatible elements(Foley et al., 2000; Gill, 1981; Klemme et al., 2005; Ryerson andWatson,1987; Schmidt et al., 2004; Schmidt et al., 2009).

Hydrous metasomatism of the mid-ocean-ridge basalt (MORB)-depleted asthenospheric mantle wedge causes partial melting bylowering the solidus of peridotite (Arculus, 1994; Kushiro et al., 1968;Stolper andNewman, 1994). This occurs either throughdirect infiltrationmetasomatism by slab-derived fluids percolating into the hot inner zoneof themantlewedge (Bourdonet al., 2003;Groveet al., 2006;Kelley et al.,2010; Peacock, 1993), or by convective corner-flow mixing of metaso-matized peridotite into these hotter central regions (Fig. 1; Schmidt andPoli, 1998; Tatsumi, 1986; Wysoczanski et al., 2006).

Partial melting of hydrated peridotite under these conditions inthe mantle wedge generates high-Mg basalts (Greene et al., 2006;Pichavant et al., 2002; Smith et al., 2010). Such arc basalts aredistinguished from MORB by higher contents of incompatibleelements (as noted above) and water (up to 6 wt.% H2O; Cervantesand Wallace, 2003; Grove et al., 2003; Pichavant et al., 2002; Sobolevand Chaussidon, 1996). Critically, Hamada and Fujii (2008) andZimmer et al. (2010) report that a water content of 2 wt.% separates“dry” tholeiitic (olivine+plagioclase/orthopyroxene) from “wet”calc-alkaline (clinopyroxene+magnetite) magmatic fractionationtrends.

Arc basalts are also characterized by distinctly higher oxidationstates than MORB (up to 2 log units above the fayalite–magnetite–quartz buffer: ΔFMQ+2; Ballhaus, 1993; Brandon and Draper, 1996;

Parkinson and Arculus, 1999; Rowe et al., 2009). The relatively highoxidation state of arc magmas is a critical factor in their subsequentmetallogeny, and originates from oxidative seafloor alteration of theoceanic plate (Staudigel et al., 1996), transmitted into themantlewedgeby the metasomatic fluid flux (Brandon and Draper, 1996; Kelley andCottrell, 2009; Malaspina et al., 2009).

2.1.1. Behavior of metalsMost base and precious metals would be expected to have at least

moderate solubilities in the hot, relatively oxidized, saline aqueousfluids exsolved from the downgoing slab. In particular, as noted inSection 2.1, Pb, As, and Sb are stronglymobilized by such fluids, possiblyalong with Tl and Cu (Noll et al., 1996). The behavior of highlysiderophile elements (HSE) such asAu andPGE is lesswell knownunderthese conditions, but studies of metasomatized mantle xenoliths fromisland arc lavas suggest that Au, Re, and the Pd-group elements(including Pt) are mobilized into the mantle wedge during subductionmetasomatism (Dale et al., 2009; Kepezhinskas et al., 2002; McInnes etal., 1999; Sun et al., 2004a; Widom et al., 2003).

The volumetric extent and efficiency of mobilization of metals intothemantlewedge by this process are unknown, butfluidmetasomatismclearly represents one viable mechanism for metal transfer into arcmagma sources.

The behavior of chalcophile and siderophile metals during subse-quent partial melting of the metasomatized mantle wedge dependscritically on oxidation state (fO2) and sulfur fugacity (fS2), because theseparameters control the stability and abundance of sulfide phases. Goldand PGE partition strongly into sulfide phases relative to silicate melts,but Cu to a somewhat lesser degree (Campbell and Naldrett, 1979;Peach et al., 1990), so if sulfides are abundant in the magma sourceregion, partial melts will be depleted in Au and PGE relative to Cu(Fig. 2). Under the high fO2 and fS2 conditions of the supra-subductionzonemantlewedge, thebulk of the sulfurfluxwill likely consist of SO2 or

5J.P. Richards / Ore Geology Reviews 40 (2011) 1–26

sulfate, dissolved first in slab fluids and then magma (Carroll andRutherford, 1985; Jenner et al., 2010; Jugo et al., 2005a). However,because of the equilibria betweenvarious sulfur species, at high fS2 somecondensed sulfide phases will likely also be present (McInnes et al.,2001). Consequently, Richards (2009) suggested that normal arcmagmas will be minimally depleted in Cu (due to its higher abundanceand moderate chalcophile nature) relative to sparse, highly siderophileelements suchasAu and PGE,whichwill be strongly retained in residualsulfide phases in the source region (e.g., Hamlyn et al., 1985; MitchellandKeays, 1981; Peachet al., 1990). Thismay explain theCu-rich natureof typical arc-related porphyry deposits (relative to Au and PGE;Richards, 2005). In contrast, Au-rich porphyry deposits may requireatypical subduction-related or collisional tectonic settings and petro-genetic processes, which act to destabilize residual sulfide phases andrender Au incompatible (e.g., Jégo et al., 2010; Richards, 1995, 2009;Sillitoe, 2000; Solomon, 1990; Wyborn and Sun, 1994; see Section 2.5).The low abundances of PGE in many arc-related ore deposits suggest afurther separation of these elements from Au and Cu, perhaps throughthe formation of residual platinoid alloy phases (e.g., Barnes et al., 1985;Borisov and Palme, 1997; Kepezhinskas et al., 2002; Peach et al., 1990)or Cr-spinels (into which Ir-group PGE strongly partition; Hattori et al.,2010; Righter et al., 2004).

2.2. Sediment dehydration and/or melting in subduction zones

Seafloor sediments on the surface of the downgoing slab areanother potential source of metasomatic contributions to the mantlewedge. Much of this sedimentary material will be scraped off at thetrench to form an accretionary prism, but varying amounts may alsobe subducted, depending on the degree of coupling between theupper and lower plates, and also the sediment input load (Fig. 1). Suchsediments will be water-rich and pelitic in bulk composition, and thusare more likely to undergo partial melting under subduction zoneconditions than basaltic oceanic crust (Hermann and Spandler, 2008).

Au

(ppb

)C

u (p

pm)

R = (mass of silicate melt)/(mass of sulfide)

Cu maximized inmagma (R ≥ 103)

Au maximized inmagma (R ≥ 106)

0.1

1

10

104

105

106

10–5

1 10 104 105 106 107 108

Cu in sulfide

Cu in magma

Au in sulfide

Au in magma

Sulfide/silicate melt partitioncoefficients:D(Cu) = 103

D(Au) = 105

Metal concentrations inmagma in absence of sulfide:Cu = 50 ppmAu = 5 ppb

Au depleted in magma

Au enriched in

residual sulfide

Porphyry Cupotential arc

magmas

Porphyry Cu-Aupotentialmagmas

102 103

102

103

10–4

10–3

10–2

5 ppb Au

50 ppm Cu

Fig. 2. Concentrations of Cu and Au in silicate magma and coexisting sulfide liquid as afunction of R=[mass of silicate melt]/[mass of sulfide melt] (Campbell and Naldrett,1979; diagram modified from Richards, 2005, 2009). At R-factors below ~102, magmaswill be depleted in both Cu and Au. At R-factors between ~102–105, magmas will bedepleted in Au but essentially undepleted in Cu (porphyry Cu-potential magmas). At R-factors N105, magmas will be undepleted in Au and Cu (porphyry Cu–Au–potentialmagmas). A corollary of this diagram is that arc magmatismwill leave small amounts ofrelatively Au-rich sulfide in the mantle source or lithosphere during fractionation,which can be remelted during post-subduction tectonomagmatic processes, to formsmall-volume, alkaline, porphyry Au-potential magmas.

Nevertheless, Aizawa et al. (1999); Dreyer et al. (2010); and Duggenet al. (2007) have suggested that dehydration is the principal processaffecting sediments down to depths of ~100 km (i.e., to below thevolcanic arc), with melting only occurring significantly at greaterdepths when temperatures exceed ~800 °C (possibly reflected in thegeochemistry of some back-arc magmas).

Sediment contributions to the source of arc magmas have been thesubject of numerous studies, with the least ambiguous evidencecoming from island arcs (e.g., MacDonald et al., 2000; Thirlwall et al.,1996; Wysoczanski et al., 2006). In continental arcs, it can be difficultto distinguish between chemical and isotopic signatures fromsubducted continent-derived sediment versus crustal contaminationduring magma ascent (e.g., Hildreth and Moorbath, 1988; Kemp et al.,2007): both sources will contribute incompatible elements andcrustal isotopic values to primary mantle-derived arc magmas(Breeding et al., 2004).

Trace elements commonly used as indicators of sediment contribu-tions to arcmagmas are Ba, B, Be, Th, andPb (Dreyer et al., 2010; Johnsonand Plank, 1999), and Ba/La and Th/La ratios can be used as ameasure ofsediment versusmantle source components (Plank, 2005;Walker et al.,2001). In particular, the cosmogenic radioisotope 10Be can be used as atracer of recent (b10 Ma) introduction of sediments into arc magmasources (Dreyer et al., 2010; Morris et al., 1990). However, although aclear sediment-derived isotopic signature can be observed in manyisland arc systems, the volumetric contribution of sediments to islandarc magmas seems to be relatively minor (Hawkesworth et al., 1994;Kilian and Behrmann, 2003; Poli and Schmidt, 2002; Stern et al., 2006).

One additional element that may be added to the mantle wedgefrom subducted sediments is sulfur, as suggested by the positive δ34Scompositions of arc magmas (de Hoog et al., 2001), which are similarto those of seafloor sediments (Alt et al., 1993). Analyses of glassinclusions in olivine from primitive arc magmas reveal concentrationsof up to 2900 ppm S (de Hoog et al., 2001), and Jugo et al. (2005b)measured experimental concentrations of up to 1.5 wt.% S in oxidizedarc basalts. These high sulfur contents have great significance for thebehavior of chalcophile and siderophile metals (see Sections 2.1.1,2.5.1, and 3).

2.2.1. Behavior of metalsLead, which is significantly enriched in the continental crust (and

crustally-derived sediments) relative to themantle, is the onlymetal forwhich a clear sedimentary source can be inferred in some island arcmagmas, because it can be readily identified by its radiogenic isotopiccomposition compared with depleted mantle sources. However, incontinental arcs, distinguishing a subducted sediment source ofradiogenic Pb from crustal contamination during magma ascent isvery difficult (e.g., Barreiro, 1984; Chiaradia et al., 2004; Kontak et al.,1990). This has led to diverging opinions: for example, Aitcheson et al.(1995);Hildreth andMoorbath (1988); James (1982); Kay et al. (1999);and Tilton et al. (1981) concluded that the bulk of the radiogenic Pb incentral Andean magmas comes from upper plate crustal sources,whereas Macfarlane (1999); McNutt et al. (1979); Mukasa et al.(1990); and Sillitoe and Hart (1984) preferred a subducted sedimentsource for Pb in some Andean igneous rocks and ore deposits.

Inferring a seafloor sediment source for othermetallic componentsin arc magmas (and related ore deposits) is much more speculative,and generally involves the subduction of metal-rich manganesenodules or even massive sulfide deposits. The latter, however, likelyoxidize and disperse geologically rapidly after formation on theseafloor (e.g., Edwards, 2004; Herzig et al., 1991) unless quicklyburied by lava; theymight thus only be expected to be subductedwithvery young oceanic crust. Few studies have specifically invoked asubducted sediment source for ore metals other than a component ofPb, and the majority of authors have concluded that such a source iseither unnecessary or unproven (e.g., Burnham, 1981; Chiaradia et al.,2004; de Hoog et al., 2001; Fontboté et al., 1990).

6 J.P. Richards / Ore Geology Reviews 40 (2011) 1–26

2.3. Oceanic slab melting in subduction zones

The question of whether the downgoing oceanic slab, or morespecifically the basaltic oceanic crust, melts during subduction hasprompted lively debate recently, not only in the petrology literature(e.g., Conrey, 2002; Defant and Kepezhinskas, 2001, 2002; Garrison andDavidson, 2003), but also amongst economic geologists because of thesuggestion that “slabmelts”might in somewaybeuniquely fertile for laterporphyry ore formation (e.g., Mungall, 2002; Oyarzun et al., 2001, 2002;Sajona and Maury, 1998; Thiéblemont et al., 1997; for contrary opinions,see: Rabbia et al., 2002; Richards, 2002; Richards and Kerrich, 2007).

The idea that the hydrated basaltic ocean crust might melt duringsubduction was an early assumption of the plate tectonic revolution,because it seemed conveniently to explain the relatively felsiccomposition of arc magmas (as opposed to basalts formed by meltingof peridotitic mantle). The theory was given credence by theexperiments of Rapp et al. (1991) and Rapp and Watson (1995), whoshowed that melting of amphibolite under upper mantle conditions(1025 °C and 0.8–1.6 GPa) could generate an intermediate compositiontonalite–trondhjemitemelt, not dissimilar to anarc andesite. Defant andDrummond (1990) termed the products of subducted slab melting“adakites”, after a single anomalous lava flow on Adak Island in theAleutians described by Kay (1978). Because garnet should be present inthe eclogitic source of these magmas, Defant and Drummond (1990)argued that such melts could be distinguished by anomalously lowconcentrations of heavy rare earth elements (HREE) and Y (which arecompatiblewithgarnet) relative to light rare earthelements (LREE), andhigh concentrations of Sr (because of the absence of plagioclase at suchdepths). Thus, slab melts, or adakites, could be distinguished on Sr/Yversus Y, or La/Yb versus Yb diagrams from normal arc magmas formedin the absence of garnet.

However, despite the theoretical possibility of melting subductedoceanic crust, most thermal models of subduction zones indicate thattemperatures in the slab do not normally reach melting conditions(N800 °C) prior to dehydration and eclogitization (Fig. 1), which wouldrender the slab infusible (e.g., Davies and Stevenson, 1992; Peacock,1996; Poli and Schmidt, 2002). Thus, Defant and Drummond (1990)proposed that slab melting might be restricted to the subduction ofyoung (≤25 m.y. old) and therefore still hot oceanic crust, and Peacocket al. (1994) were even more restrictive (b5 m.y. old). Other relativelyuncommon scenarios that might result in slab melting include shallowor stalled subduction (whereby the slab has more time to heat up atshallow depths; Gutscher et al., 2000; Peacock et al., 1994), ridgesubduction (Guivel et al., 2003; Kay et al., 1993), or edge-melting ofdetached slabs or slab windows (Haschke and Ben-Avraham, 2005;Thorkelson and Breitsprecher, 2005; Yogodzinski et al., 2001).

A further complication is added by the fact that most adakitesdescribed in the petrology literature are not in fact primary slabmelts,but are substantially evolved, having reacted or hybridized with theasthenosphere during ascent (and likely also the upper platelithosphere). This modification to the adakite slab-melting model isrequired to explain the high contents of MgO, Ni, and Cr present insome adakites relative to expected levels for hydrated basalt partialmelts (Defant and Kepezhinskas, 2001; Drummond et al., 1996;Martin, 1999; Martin et al., 2005; Yogodzinski et al., 1995).

Direct evidence for slabmelting is lacking, but supra-subduction zonexenoliths from the Tabar-Lihir-Tanga-Feni (Papua New Guinea), Philip-pine, and Patagonian arcs preserve hydrous, silica-rich glass inclusionsthat are thought to represent migrating slab melts (respectively: Kilianand Stern, 2002; McInnes and Cameron, 1994; Schiano et al., 1995). Theglass inclusions characterized by Schiano et al. (1995) were calc-alkalinein composition, with high incompatible element and low Ti, Nb, and Ycontents, high LREE/HREE ratios, and homogenization temperatures of~920 °C. They thus compositionally resemble melts that would bepredicted to form from slabmelting, and appear to provide evidence thatthis process occurs at least locally where conditions permit.

The chemical difference between slab dehydration and slabmelting would seem to be rather small, given that both mediawould be enriched in volatiles, incompatible elements, and silica.Indeed, as noted in Section 2.1, there may well be a continuumbetween silica-rich aqueous fluids and aqueous silicate melts atgreater depths in subduction zones (Kawamoto, 2006; Kessel et al.,2005a,b; Manning, 2004; Portnyagin et al., 2007). This likely explainswhy the debate between slab melting and dehydration is somewhatintractable, and mostly hinges on subtle trace element characteristics.

2.3.1. Behavior of metalsSlab melts are predicted to be volatile-rich (including H2O, S, and

Cl) and oxidized, and thus, like hydrous slab fluids, would be expectedto be able to transport base and precious metals at least to somedegree. However, analyses of such metals (except iron) are notreported in most melt inclusion studies (e.g., Kilian and Stern, 2002;McInnes and Cameron, 1994; Schiano et al., 1995), so there are nodirect constraints on the capacity of suchmelts to transfer chalcophileand siderophile metals from the slab to the mantle wedge.

In a study of metasomatized mantle xenoliths from a submarinevolcano near Lihir Island, Papua New Guinea, McInnes et al. (1999)concluded that enrichments in Cu, Au, and PGEwere caused by slab fluidmetasomatism, rather thanmelts. In contrast, Kepezhinskas et al. (2002)measured the concentrations of Au and PGEs in mantle xenoliths fromthe Kamchatka arc, and suggested that a fluid-transported componentcould be distinguished froma slabmelt component by co-enrichments inPGE and high field strength elements (HFSE) in the latter, because of thelow capacity of aqueous fluids to carry HFSE. Intriguingly, they noted nosuch correlation betweenHFSE-enrichments andAu, and concluded that,whereas PGEmight be transported into themantle wedge by both fluidsand melts, Au was likely only carried by hydrous fluids.

Two theoretical studies have proposed that slab melts should beunusually effective as metal-transporting and ore-forming agents.Oyarzun et al. (2001) argued that slab melts should be unusuallyoxidized and rich in H2O and SO2 (relative to normal arc magmasderived by asthenospheric partial melting), although no evidence wasgiven for this assertion. Such magmas, they argued, should beparticularly suitable for the formation of magmatic–hydrothermalporphyry copper deposits upon emplacement in the upper crust.

Mungall (2002) presented a theoretical model for oxidation of themantle wedge by Fe3+-rich slab melts to the point of complete sulfidedestruction, thereby rendering chalcophile and siderophile elementsincompatible in mantle phases, and free to partition into silicatemelts. He argued that ferric iron is a much stronger oxidant than slab-derivedwater, and that slab melts should be rich in Fe3+ generated byoxidative seafloor alteration. Thus, slab melts might be uniquelyfavorable for the subsequent generation of metal-rich, and particu-larly Au-rich, magmas derived from the mantle wedge.

Mungall's (2002)modelmay have applicability for less commonAu-rich porphyry deposits formed in atypical subduction zone settings thatmight cause slab melting, but does not seem well suited to explainregular arc porphyry Cu deposits. In either case, metals are envisaged tobe sourced from themantle wedge, not the slab. In contrast, Oyarzun etal.'s (2001) model does not address the source of metals, and is at rootbased on the assumption that slab melts are uniquely more H2O- andSO2-rich, and more oxidized than normal arc magmas, leading tospecific ore depositional processes rather than source processes. It is notclear that these assumptions are justified, and some have argued thatslabmelts might in fact be relatively reducing because of the additionalpresence of organic-rich sediment melts (Wang et al., 2007a).

2.4. Supra-subduction zone lithospheric melting: the MASH process

Hydrous basaltic magmas generated in themantlewedgewill havetemperatures in excess of 1000 °C (Eggins, 1993; Grove et al., 2006;MacDonald et al., 2000), and perhaps as high as 1350 °C (Schmidt and

7J.P. Richards / Ore Geology Reviews 40 (2011) 1–26

Poli, 1998; Tatsumi, 2003). Because their densities will be lower thanthe mantle but not the crust (Herzberg et al., 1983), they will tend torise from their asthenospheric source region and penetrate themantlelithosphere, but pool at the crust/mantle density barrier (“level ofneutral buoyancy”: Fig. 3; Fyfe, 1992; Hildreth, 1981). Here, if the fluxof magma and heat is maintained and supplemented by the latentheat of crystallization as the magma begins to crystallize, hightemperatures can be brought to bear on lower crustal rocks that willcause partial melting (Annen et al., 2006; Bergantz and Dawes, 1994;Huppert and Sparks, 1988; Klepeis et al., 2003; Petford and Gallagher,2001; Rushmer, 1993). Hildreth and Moorbath (1988) suggested thatit is the interaction between this hot, hydrous basalt flux from thesubduction zone and felsic crustal partial melts that gives rise to theuniform composition of andesites in continental volcanic arcs, by aprocess they dubbed melting–assimilation–storage–homogenization(MASH). In a refinement of this model, Annen et al. (2006) referred tothe region of magma–crust interaction as a “hot zone” (Fig. 3).Because garnet is a product of such lower crustal fractionation andpartial melting processes (Alonso-Perez et al., 2009; Berger et al.,2009; Dufek and Bergantz, 2005; Garrido et al., 2006; Hansen et al.,2002; Klepeis et al., 2003; Rushmer, 1993; Wolf and Wyllie, 1994),derivative calc-alkaline magmas may display trace element compo-sitions that resemble adakites (see Section 2.3) but which areunrelated to slab melting (Klepeis et al., 2003; Macpherson et al.,

Volcanic arc

Sea level

1000°C

Oceanicmantle

lithosphere

M

Feeder

600°C

Dehydrating oceanic crust

EHydrothermalalteration P

Fig. 3. Schematic section through a continental arc, showing the development of a MASH or “hbuoyancy, differentiate, and interact with crustal rocks and melts. Evolved, less dense, andesitbuoyancy to form batholithic complexes. Along with volcanic structures, porphyry and epithexsolvedmagmatic fluids ascend, cool, and interact with near-surface upper crustal rocks. ModAnnen et al. (2006), and Sillitoe (2010).

2006; Richards and Kerrich, 2007; Tulloch and Kimbrough, 2003).Such common processes, affecting batches of magma crystallizing andfractionating at different crustal depths (e.g., Annen et al., 2006) areentirely consistent with petrological observations in arc volcanicsystems where “adakite-like” (i.e., high-Sr/Y) andesitic lavas may beinterlayered with “normal” andesites in a single volcano, and do notrequire a fundamental change in magma source 100 km below thevolcano (e.g., Feeley and Davidson, 1994; Grunder et al., 2008;Richards et al., 2006a).

Once these hybrid magmas reach basaltic andesitic to andesiticcompositions, their densities are low enough to allow them to risethrough the lower continental crust (Herzberg et al., 1983), but theytend to stall again at a second density barrier in the middle to uppercrust below light supracrustal rocks. This is the level (5–10 km) atwhichlarge arc batholiths will form if the flux of magma is sustained, and isalso the level atwhichevolved felsicmelts andvolatiles are accumulated(Fig. 3; see Richards, 2003, and references therein). These volatiles drivebuoyant, bubbly, evolved magma upwards into the cover rocks to formsubvolcanic stocks and dikes, or explosive volcanic eruptions if theyreach surface. The volatiles may also separate from the magma flux toform a separate fluid plume, which ultimately vents at the surface(fumaroles) butmay also formporphyry- and epithermal-type depositsin the hypabyssal and near-surface environment (see discussion ofthese processes in Sections 3 and 4).

1400°C

600°C

1000°C

Continental crust

Metasomatizedasthenosphere

1400°C

Subcontinentalmantle lithosphereantle corner flow

Asthenosphere

0 km

50 km

100 km

Mid- to upper crustalbatholithic complex

Lower crustalMASH or “hot zone”

dikes

pithermal depositsorphyry deposits

Partial melting

ot zone” at the base of the crust where basaltic arc magmas pool at their level of neutralic magmas rise into the mid-to-upper crust where they pool at their new level of neutralermal deposits may form at shallower levels above these batholithic complexes whereified from Richards (2003, 2005); sources: Hildreth andMoorbath (1988), Winter (2001),

8 J.P. Richards / Ore Geology Reviews 40 (2011) 1–26

2.4.1. Behavior of metalsSome of the clearest evidence for the involvement of crustal rocks in

continental arcmagmas comes fromPb isotopic data (e.g.,Wörner et al.,1992), although as noted in Section 2.2, there may be ambiguitybetween lower crustalmelting and themelting of subducted continent-derived sediments. Lead and uranium are much more abundant in thebulk continental crust (11 ppm Pb, 1.3 ppm U; Rudnick and Gao, 2003)or lower continental crust (4 ppm Pb, 0.2 ppm U; Rudnick and Gao,2003) than the primitive mantle (b0.2 ppm Pb, b0.02 ppm U; Taylorand McLennan, 1985; Sun and McDonough, 1989), MORB (0.3 ppm Pb,0.05 ppm U; Sun and McDonough, 1989), or typical low-K mafic arcandesites (b1.8 ppm Pb, b0.2 ppm U; Gill, 1981), so it only requiressmall amounts of contamination by radiogenic crustal lead tosignificantly modify a magma's Pb isotopic composition. Thus, it is notclear that a particularly large amount of Pb in arc magmas is sourcedfrom crustal rocks versus the subduction zone, and Macfarlane et al.(1990) have argued that the crustal contribution is minimal in CentralAndean magmas and ores. Moreover, porphyry-type systems aretypically not Pb-rich, except in late-stage skarns and distal veinswhere some of the Pbmay have been derived from local host rocks (e.g.,Mukasa et al., 1990).

Most researchers have assumed that, because of the higherconcentrations of Cu in primitive andesites (145 ppm; Gill, 1981)compared with the bulk continental crust (27 ppm; Rudnick and Gao,2003), the bulk of Cu in porphyry-type deposits is mantle-derived. Asimilar assumption is made for Au, although the primitive mantle andcontinental crust actually have comparable concentrations (1–3 ppbAu; Rudnick and Gao, 2003; Taylor and McLennan, 1985). Moreover,porphyry Cu–(Au) deposits are found in association with mantle-derived arc magmas worldwide, regardless of crustal type (oceanic orcontinental) or thickness (Kesler, 1973), so a crustal heritage for thesemetals does not appear to be critical. Nevertheless, a lower crustalsource, perhaps hybridized with mantle-derived magmas, has beenproposed by Bouse et al. (1999) for both magmas and metals in theLaramide porphyry systems of Arizona. Moreover, Titley (1987, 2001)has specifically suggested that Au and Ag are crustally derived in arange of ore deposits including porphyries in southwestern USA,because the ratios of these elements correlate closely with twodistinct basement domains in this region. Given that Au is notespecially enriched in the mantle (see above), and that Ag is in factmore abundant in the crust than the mantle (80 ppb in the bulkcontinental crust, versus b19 ppb in the primitive mantle; Taylor andMcLennan, 1985), a crustal source for at least some proportion ofthese minor metals, and especially Ag, in arc magma-related systemsmay be reasonable. However, it seems unlikely that this argument canbe extended to copper, except perhaps on the margins.

2.4.2. Sources of MoMolybdenum occurs in varying amounts in porphyry-type deposits,

ranging from trace levels (b0.01 wt.%Mo) in porphyry Cu–(Mo)deposits,where itmay not even be recovered as a byproduct, to being themain orecomponent (up to 0.3 wt.% Mo) in porphyry Mo deposits (Seedorff et al.,2005;Westra and Keith, 1981). At theMo-rich end of the spectrum, thereare two clearly different tectonomagmatic associations, only one ofwhichis directly related to subduction: calc-alkaline porphyry Mo deposits aregenerally relatively low grade (0.1–0.02 wt.% Mo; Carten et al., 1993),whereas intra-cratonic rift-related deposits associated with high-silica,fluorine-rich, peraluminous granitoids are relatively high grade (0.1–0.3 wt.%Mo;e.g., “Climax-type”deposits; Cartenet al., 1993;KirkhamandSinclair, 1996; Sinclair, 2007; Stein, 1988; White et al., 1981).

Kesler (1973) noted a general association (with exceptions) ofporphyry Cu–Au deposits in island arcs, and porphyry Cu–Mo depositsin continental arcs, and it is clear that the peraluminous felsic rocksassociated with rift-related Climax-type porphyry Mo deposits areprimarily of continental crustal origin (Farmer and DePaolo, 1984;Stein, 1988). This has led to one view that Mo might be predominantly

derived from continental crustal sources (Farmer and DePaolo, 1984;Stein, 1988; Klemm et al., 2008; White et al., 1981). On the other hand,minor amounts of Mo do occur in some island arc-related porphyrydeposits where no continental crustal sources are inferred (Westra andKeith, 1981), so amantle (subduction zone) source for at least someMocannot be excluded. Moreover, Blevin and Chappell (1992) and Blevinet al. (1996) have demonstrated a continuum from Cu–Au depositsassociated with unevolved, mafic I-type granitoids to W–Mo depositsassociated with cogenetic, evolved granites in eastern Australia,suggesting a common, magmatic source for all of these elements.

A complication is introduced in the Climax-type deposits, becausealthough the immediate source of the Mo-bearing fluids is felsicmagma of clear crustal origin, many deposits also show a close geneticassociation with mafic alkaline magmas, which may have introducedvolatiles, S, and possibly Mo into the evolved felsic magma chamber(Audétat, 2010; Carten et al., 1993; Keith et al., 1986, 1998). Keith etal. (1997), Hattori and Keith (2001), and Maughan et al. (2002) havealso suggested that injections of mafic alkaline magmas into theevolving Bingham Canyon magmatic system may have given rise tothe unusually large size and high grades of this porphyry Cu–Mo–Audeposit. Along the same lines, Pettke et al. (2010) have proposed thatthe unusual Cu–Mo–Au endowment of the southwestern USA (e.g.,the giant Bingham, Butte, Climax, Henderson, and Questa porphyryCu–Mo–Au and porphyry–Mo deposits) reflects Cenozoic remobiliza-tion of Proterozoic subduction-metasomatized subcontinental mantlelithosphere (see Section 2.5).

Thus, at this time there is no consensus regarding the crustalversusmantle origin ofmolybdenum in porphyry deposits, although itis clear that the highest grade porphyry Mo deposits are formed inintra-plate continental settings, and if a mantle source is important inthese cases, it is not directly related to subduction activity but ratherto rifting or reactivation of previously subduction-enriched litho-spheric sources.

2.5. Lithospheric melting during post-subduction events

A number of mineral deposits with broad similarities to thoseformed by subduction-related processes are also found in post-subduction tectonic settings, such as subduction reversal or migration,arc collision, continent–continent collision, and post-collisional rifting.They include porphyry Sn–W, Mo, Cu–Mo, and Cu–Au deposits andepithermal Au deposits, and in many cases are only known not to bedirectly related to subduction because of precise geochronology andplate tectonic reconstructions that place their formation after subduc-tion has demonstrably ceased. Associated magmas are typically calc-alkaline, but tend towards somewhat more alkaline compositionscompared with normal arc magmas (Richards, 2009).

In complex accretionary arcs, it can be very difficult to ascribe anygiven pluton (and any associated mineral deposits) to a particularsubduction or collisional event, because subduction commonly con-tinues after collision, albeit normally with a shift in the locus ofmagmatism. However, in continent–continent collision zones or wherearc collision terminates subduction, there can be greater certainty aboutthe timing of cessation of subduction magmatism. Consequently, it is incollisional orogens such as theNeo-Tethyan belt of southeastern Europeand southernAsia that someof the clearest examples of post-subductionmagmatism and mineralization are found. These include, from east towest, theMioceneGangdese porphyry Cu–Mobelt in the Tibetanorogen(Hou et al., 2006, 2009; Yanget al., 2009), theMioceneKermanporphyryCu–Mo belt in southeastern Iran (Shafiei et al., 2009), the Miocene SariGunay epithermal Au deposit in northwestern Iran (Richards et al.,2006b), theEoceneÇöpler epithermalAudeposit in southeasternTurkey(Keskin et al., 2008; Kuscu et al., 2010), the Pliocene Kisladag porphyryAu deposit in western Turkey, the Miocene Skouries porphyry Cu–Au–PGEdeposit inGreece (Economou-Eliopoulos andEliopoulos, 2000), and

9J.P. Richards / Ore Geology Reviews 40 (2011) 1–26

the Roşia Montană epithermal Au deposit in Romania (Manske et al.,2006; Neubauer et al., 2005).

Similarly, in the southwest Pacific ocean, accurate plate tectonicreconstructions permit the identification of a number of post-subduction porphyry and epithermal deposits, such as the Grasbergporphyry Cu–Au deposit in Papua, Indonesia (Cloos et al., 2005;Paterson and Cloos, 2005), the Ok Tedi porphyry Cu–Au deposit (vanDongen et al., 2010) and the Porgera alkalic-type epithermal Au depositin mainland Papua New Guinea (Richards et al., 1990; Richards andKerrich, 1993), the Lihir alkalic-type epithermal Au deposit on LihirIsland, PapuaNewGuinea (Carman, 2003;Kennedyet al., 1990), and theEmperor alkalic-type epithermal Au deposit in Fiji (Gill and Whelan,1989; Setterfield et al., 1992). (For reviews of alkalic-type epithermaldeposits, see Jensen and Barton, 2000 and Richards, 1995).

Because subduction has ceased in these regions, a fresh supply offluids, volatiles, and other slab-derived components to the mantlewedge no longer exists. Nevertheless, the broad geochemical similarityof many of these magmas to normal arc magmas, including theirhydrous and generally oxidized nature, suggests some link tosubduction metasomatism. Consequently, many researchers haveimplicated upper-plate lithospheric sources, modified by earliersubduction-related fluids and/or hydrous melts (e.g., Clemens et al.,2009; Cloos et al., 2005; Guo et al., 2007; Harris et al., 1986; Johnsonet al., 1978; Pearce et al., 1990; Pettke et al., 2010; Richards, 2009).Previously subduction-modified asthenosphere is unlikely to be a viablesource except for a short period after subduction has ceased (e.g.,Richards et al., 1990; Solomon, 1990), because such material will bequickly dispersed by mantle convection.

The key to all of these models is subduction-derived water, which ismost likely stored in amphibolitic cumulates, residual from the earlierarcmagmaflux and located in thedeep crust ormantle lithosphere (e.g.,Claeson and Meurer, 2004; Davidson et al., 2007; DeBari and Coleman,1989; Jagoutz et al., 2009; Larocque and Canil, 2010; Müntener andUlmer, 2006; Tiepolo and Tribuzio, 2008). Water lowers the solidus ofsilicate assemblages, and will lead to the formation of hydrous partialmelts during pro-grade metamorphism or mafic melt invasion (Beardand Lofgren, 1991; Rushmer, 1991; Wolf and Wyllie, 1994).

Thermal rebound in thickened orogenic crust, delamination ofsub-continental mantle lithosphere, or post-collisional rifting (withingress of asthenospheric melts into the lower crust in the last twocases) can all cause small-volume partial melting of arc-metasoma-tized lithosphere and/or hydrous lower crustal cumulates (Fig. 4;Brown, 2010; Clemens et al., 2009; Harris et al., 1986; Richards, 2009).Such melts, being derived from subduction-modified sources, willshare many of the characteristic geochemical features of arc magmas,including their relatively high water contents and oxidation states,and potentiallymetal contents (see Section 2.5.1). The smaller volumeof partial melting to be expected in such tectonic settings will givethese magmas a somewhat more alkaline composition than arcmagmas (e.g., Clemens et al., 2009), and will also mean that largebatholithic complexes are unlikely to be formed, consistent with thegenerally smaller and more isolated occurrence of such post-subduction magmatic systems (compared with arc-related Cordille-ran batholiths, or collisional S-type batholiths; Pitcher, 1997).

Because these post-subduction magmas are derived from amphibo-litic sources in which garnet (±titanite) is likely also present, andbecause their hydrous nature will suppress plagioclase fractionation(similar to other hydrous arc magmas), they may be characterized byelevated Sr/Y and La/Yb ratios; that is, they may display adakite-liketrace element characteristics.

2.5.1. Behavior of metals in subduction-modified sourcesAs discussed in Section 2.1, arc magmas are characterized by high

fO2 and fS2 relative to normal melts from MORB-depleted astheno-sphere. Consequently, such magmas may be sulfide-saturated (atoxidation states up to ΔFMQ+2.3; Jugo, 2009) despite sulfur being

predominantly present in the melt as sulfate or SO2 (Carroll andRutherford, 1985). Xenoliths from supra-subduction zone mantle(McInnes et al., 1999) and samples of mafic cumulates from lowercrustal arc roots (Fig. 5; Canil et al., 2010; Greene et al., 2006; Jagoutzet al., 2007) reveal the common presence of small amounts of sulfide,typically trapped as inclusions in silicate phases (suggesting a primarymagmatic rather than secondary hydrothermal origin).

Hamlyn et al. (1985), Richards (1995, 2009), Solomon (1990), andWyborn and Sun (1994) have explored the role of residual sulfidephases on the metal content of fractionating magmas, and also ofpartial melts formed during later, post-subduction melting events.The high partition coefficients for chalcophile and highly siderophileelements (HSE) between sulfide phases and silicate melt mean thatsuch metals should be strongly partitioned into any coexisting sulfidephases (Campbell and Naldrett, 1979; Peach et al., 1990). As shown inFig. 2, at high abundances of sulfide relative to silicate melt (low R-factor; Campbell and Naldrett, 1979), the melts will be depleted in allof these chalcophile and siderophile elements. In contrast, atintermediate abundances of sulfide (intermediate R-factor), onlyoriginally sparse HSE will show significant depletions. This ledRichards (2005, 2009) to propose that small amounts of sulfide leftbehind as residual phases from fractionation of arc magmas in thedeep lithosphere (or asthenosphere) will not significantly depletethose magmas in relatively abundant chalcophile elements such as Cuand Mo, but might significantly deplete them in highly siderophileelements such as Au. This would give rise to magmas with relativelyhigh Cu/Au ratios (which might form Cu-rich porphyry deposits), butwould leave a residue of potentially HSE-rich sulfides in the mantleand/or lower crustal amphibolitic cumulate arc roots.

As noted in Section2.5, subduction-modified asthenospheric sourceswill be rapidly convected away when subduction ceases, and so couldonly contribute to immediately post-subduction magmatism. Incontrast, deep crustal amphibolites are preserved in the lithosphere,andwill be susceptible to partialmeltingat any later timedue to thermalreboundor reheating by invading asthenosphericmelts. Under lower fS2post-subduction conditions (a flux of S from the subduction zone is nolonger present), any residual sulfide phases would likely dissolve intothe S-undersaturated silicatemelt, carrying their metal loadswith them(e.g., Ackerman et al., 2009). Richards (2009) proposed that this mightexplain the occurrence of Au-rich porphyry and related epithermalsystems in some post-subduction settings, such as the alkalic-typeepithermalAu deposits of the SWPacific, and various post-collisional Audeposits in the Balkans–Turkey–Iran Neotethyan belt. Pettke et al.(2010) have proposed a similar model for giant porphyry Cu–Mo–Audeposits in the southwestern USA.

This model can also explain the occurrence of Au–poor porphyryCu–(Mo) deposits in post-subduction settings, such as those in Tibetand Iran (Hou et al., 2009; Shafiei et al., 2009; Wang et al., 2007b), theonly difference being that in this case larger proportions of sulfidemay have fractionated out from the original arc magmas in the deepcrust. Such sulfides would have retained significant amounts of thesubduction flux of Cu and Mo, but HSE would be diluted to lowconcentrations by the greater volume of sulfide (low R-factor; Fig. 2).Second-stage melts from such cumulate sources would therefore beCu–(Mo)-rich, but not necessarily Au-rich.

A control on these two scenarios (abundant Cu-rich residual sulfideversus sparse but HSE-rich residual sulfide, or low versus high R-factor)might be the average oxidation state and sulfur fugacity of the generativesubduction system. Inmore oxidized or S-poor systems, smaller volumesof HSE-rich sulfide would exsolve from the silicate melt (high R-factor;Campbell andNaldrett, 1979),whereas in less oxidized or S-rich systems,larger volumes of Cu-rich but HSE-poor sulfide would exsolve (low R-factor; Fig. 2). In particular, the proportion of sulfides exsolving from arcmagmas may be very sensitive to small changes in their oxidation state,because of the rapid change from sulfide to sulfate dominance inmagmatic systems between ΔFMQ+1 and +2 (Jugo et al., 2010). The

Fig. 4. Post-subduction tectonic environments conducive to the formation of porphyry and epithermal deposits by remobilization of previously subduction-modified lithosphere(modified from Richards, 2009). (a) Porphyry Cu±Mo deposits formed in normal arc settings; a continental arc is shown, but similar processes can occur inmature island arcs. (b–d)During post-subduction tectonic processes, previously subduction-modified sub-continental lithospheric mantle (SCLM) or lower crustal hydrous cumulate zones residual fromprevious arc magmatism (black layer) may undergo small-volume partial melting. Such magmas may remobilize Au as well as Cu±Mo left behind in residual sulfide phases by arcmagmatism, leading to the potential formation of porphyry Cu±Au±Mo and alkalic-type epithermal Au deposits. Magmas may be characterized by high Sr/Y and La/Yb ratios dueto the presence of hornblende (±garnet, titanite) in the amphibolitic lower crustal source rocks. See text for discussion.

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oxidation state of the mantle wedge will depend on the character of theflux from the subducting slab (e.g., a higher proportion of subductedorganic-rich sediment would lead to lower oxidation states; Wang et al.,2007a); this property is therefore likely to have a characteristic averagevalue along any given arc at any particular period of time.

Variations in oxidation state over typical ranges for arc magmas(ΔFMQ=0 to +2; Ballhaus, 1993; Brandon and Draper, 1996; Blatterand Carmichael, 1998; Malaspina et al., 2009; Parkinson and Arculus,1999; Rowe et al., 2009) will not greatly affect the potential to form syn-subduction porphyry Cu–(Mo) deposits, but might control the Cu/HSEratio in later magmas formed by post-subduction melting of thesesulfide-bearing residues. Specifically, Au-rich (low Cu/Au) post-subduc-tion porphyries might form in settings where previous arc magmatismwas relatively oxidized (sparse but HSE-rich sulfide residue), whereasAu-poor (high Cu/Au) post-subduction porphyries might form whereprevious arcmagmatismwas relatively reduced (more abundant Cu-richsulfide residue). Such a mechanism might also explain why coeval beltsof porphyry deposits tend to have characteristic Cu/Au ratios.

Finally, partial melting of predominantly reduced, sulfide-rich crustalrocks in orogenic settings may lead to chalcophile and siderophileelement-depleted, but potentially lithophile element-rich S-typemagmas (see Section 2.6).

2.6. Crustal melting during post-collisional stress relaxation

Collisional orogens commonly undergo crustal thickening followedby extensional or transpressional collapse. Bimodal magmatism is

characteristic of such tectonic settings, resulting from partial melting ofpelitic protoliths in the deep crust triggered by the heat from upwellingasthenospheric melts (Hildreth, 1981). Peraluminous S-type granites(Chappell and White, 1974) are subsequently emplaced as largebatholith complexes in the mid- to upper orogenic crust (e.g., theHercynian peraluminous granites of Europe; Barbarin, 1996; Clemens,2003; Darbyshire and Shepherd, 1994; Harris et al., 1986; Wyllie et al.,1976). These granites tend to be enriched in lithophile rather thanchalcophile elements, reflecting their crustal origins, and may generatemagmatic–hydrothermal deposits containingSn,W,U,Mo, REE, Li, Be, B,and F.

2.6.1. Sources of metalsTin and especially tungsten commonly accompany molybdenum in

porphyry deposits as trace metals and byproducts, but they also form aclass of porphyries on their own, associated with S-type granites incontinental orogens (Hart et al., 2005; Ishihara, 1981; Ishihara andMurakami, 2006; Kerrich and Beckinsale, 1988; Kirkham and Sinclair,1996; Lehmann, 1982). The Hercynian tin granites of Europe, and theBolivian and SE Asian tin belts are examples of such deposits, withmineralization occurring in skarns and greisens around the graniteintrusions, and to a lesser extent as internal stockworks anddisseminations (e.g., Černý et al., 2005; Meinert et al., 2005). As withthe source magmas, metals in these deposits appear to be predomi-nantly of crustal origin (Hedenquist and Lowenstern, 1994). Forexample, in a recent assessment of the source of Sn in the Cornubianbatholith of SW England, Williamson et al. (2010) concluded that all of

Chalcopyrite

Chalcopyrite

Pyrite

Pyrrhotite

(a)

(b)

Fig. 5. Reflected light photomicrographs of sulfide inclusions in amphibole-rich lowercrustal arc cumulates from: (a) the Talkeetna arc, Alaska; (b) the Bonanza arc,Vancouver Island, Canada (samples courtesy A. Greene and D. Canil, respectively).

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the Sn could have been extracted from the crustally-derived granites.Uranium is also significantly enriched in crustal rocks versus themantle(0.91 ppm versus ~0.02 ppm, respectively; Taylor and McLennan,1985), and so is unlikely to have a mantle source in such deposits.However, Dietrich et al. (1999) have suggested a possible role formantle-derived magmas in triggering volatile (andmetal) release fromevolved, felsic magmas in the Bolivian tin belt, andWalshe et al. (2011)have identified a mantle Nd isotopic signature in tin granites fromeastern Australia.

In contrast, in the case of W skarns associated with Mo mineraliza-tion in calc-alkaline I-type magmas, a shared mantle origin with Momight be indicated (e.g., Newberry and Swanson, 1986), consistentwiththe similar siderophile tendencies of these two elements, and theirposition in the periodic table (group VIB).

3. Behavior of metals during magma fractionation and fluidexsolution in the upper crust

Key to the formation of magmatic–hydrothermal deposits ofchalcophile and siderophile elements in the upper crust is the lack ofsignificant saturation with and loss of sulfide phases prior to aqueousvolatile exsolution from a cooling magma (Candela, 1989b, 1992;Candela and Holland, 1986; Candela and Piccoli, 2005; Richards, 1995;Richards and Kerrich, 1993; Spooner, 1993). As discussed in Sections2.1.1 and 2.5.1, chalcophile and siderophile elements partition

strongly into sulfide phases exsolving or crystallizing from silicatemelts. Thus, if extensive fractionation and removal of magmaticsulfide phases were to occur, the remaining silicate melt would bestrongly depleted in these elements (Jugo et al., 1999; Lynton et al.,1993). This is in essence the model for formation of orthomagmaticsulfide deposits from relatively reduced mafic magmas (e.g., Naldrett,1989), and perhaps explains the deficiency of chalcophile elementssuch as Cu and Au in more reduced, evolved, Sn–W-bearing S-typemagmas (Blevin and Chappell, 1992; Hedenquist and Lowenstern,1994).

In more oxidized subduction-related or post-subduction magmas,although sulfide saturation likely occurs at some point during theirevolution (as indicated by the common presence of sparse sulfideinclusions in phenocrysts in arc volcanic rocks as well as lower crustalarc cumulates; e.g., Burnham, 1979; Halter et al., 2002; Hattori, 1997;Keith et al., 1997; Stavast et al., 2006; Fig. 5), it may not occur to theextent that sulfides physically separate from the magma, or at leastnot in large amounts. Instead, small sulfide droplets or crystals may beentrained in magma ascending buoyantly through the crust (e.g.,Bockrath et al., 2004; Tomkins andMavrogenes, 2003) or as inclusionsin silicate phenocrysts, and will not be substantially lost to the overallmagma flux. Indeed, several authors have argued that pre-concen-tration of ore metals in magmatic sulfide phases may be an importantstep in porphyry metallogenesis (e.g., Jenner et al., 2010). In thesemodels, sulfide phases subsequently break down due to changes inoxidation state and sulfur fugacity in response to volatile exsolutionand magnetite crystallization upon emplacement in the upper crust,thereby rendering metals available for redissolution in the volatilephase (e.g., Cygan and Candela, 1995; Halter et al., 2002, 2005; Jugo etal., 1999; Keith et al., 1997; Stavast et al., 2006). Other authors haveargued that this process, while it may occur, is not critical to metalbehavior in magmatic–hydrothermal systems, and that direct parti-tioning from the silicate melt to the hydrothermal fluid phase is thedominant mechanism (e.g., Audétat and Pettke, 2006; Lynton et al.,1993; Simon et al., 2008; Sun et al., 2004b). As noted in Section 2.5.1,Richards (2009) suggested that separation of small amounts of sulfidefrom arcmagmas at depth in lower crustal MASH cumulate zonesmayprovide a source of metals (especially HSE) for later post-subductionmagmas, but that this process may not substantially affect the Cucontent of the original arc magmas.

Regardless of the exact role of magmatic sulfides, the ultimaterelationship is that of partitioning of metals between silicate melts andexsolving hydrothermal fluids, with sulfides as a possible intermediarystep. Our current understanding of these partitioning processes is nowquite advanced following several decades of hydrothermal experiments(e.g., Candela and Piccoli, 1995) and more recently the advance ofquantitative single fluid andmelt inclusion analysis (e.g., Heinrich et al.,2003a,b), which has enabled direct measurement of metal contents inore-forming fluids andmelts. In the following sections, I review some ofthe key factors in metal solvation and transport in magmatic–hydrothermal fluids.

3.1. Partitioning of metals from magma into exsolving hydrothermalfluid

Subduction-relatedmagmas commonly contain at least 4 wt.% H2Oduring crustal ascent, as evidenced by the presence of hornblende andbiotite phenocrysts in many andesitic volcanic rocks and arc plutons(Burnham, 1979; Naney, 1983; Rutherford and Devine, 1988). Inresponse to the decreasing solubility of water in silicate melts aspressure decreases, such hydrous magmas inevitably exsolve anaqueous volatile phase upon emplacement at shallow crustal levels oron eruption (Burnham, 1979, 1997; Eichelberger, 1995). This processhas in the past rather confusingly been called first and second boiling(although neither process is technically “boiling”), the first eventoccurring during ascent and depressurization of the magma, and the

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Fig. 6. P–T–XNaCl phase diagram,modified fromDriesner and Heinrich (2007), illustratingfluid pathways for amagmatic–hydrothermal fluid exsolvedwith an initial bulk salinity of10 wt.% NaCl. (a) Early, high thermal gradient fluids exsolved from deeply (path 1) andshallowly (path 2) emplaced magmas (porphyry environment). (b) Late, low thermalgradient fluids exsolved from deeply emplaced magma (high-sulfidation epithermalenvironment). Path 3 illustrates the supercritical fluid contraction path proposed byHedenquist et al. (1998), whereas path 4 illustrates a slightly steeper thermal gradientwith brief intersection of the 2-phase (L–V) field, as proposed by Heinrich et al. (2004).Note that in the two-phasefield, thedense saline liquidphase progressively separates fromthe vapour phase, and the two-phase pathways shown do not represent a closed system.See text for discussion.

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second occurring after emplacement as crystallization progressivelyincreases the concentration of volatiles in the residual melt (Candela,1989a). In reality, volatile exsolution is probably a more-or-lesscontinuous process during arc magma ascent and cooling, starting atdepth with the exsolution of relatively insoluble CO2 (Blundy et al.,2010; Holloway, 1976; Lowenstern, 2001; Wallace, 2005). However,the bulk of the magmatic water content likely exsolves relativelyrapidly as the magma approaches its solvus, at depths of ~5–10 km,depending on magma composition and water content (Burnham,1979).

The composition of this exsolved magmatic volatile phase isdominantly aqueous, containing sulfur species (predominantly SO2 inoxidized systems, but also some reduced S species), CO2, NaCl, KCl, HCl,andmetal chlorides. The exact composition depends onmany variables,including the depth of exsolution and the magma composition(especially the initialmagmatic Cl/H2O ratio and alkali content; Candela,1989c; Candela and Piccoli, 2005; Cline and Bodnar, 1991; Webster,1992), but typical estimates for a single-phase supercritical fluidexsolved at depths below the H2O–NaCl solvus are ~2–13 wt.% NaClequivalent (average 5 wt.% NaCl equivalent) with minor CO2 (Audétatand Pettke, 2003; Audétat et al., 2008; Burnham, 1979; Candela, 1989c;Hedenquist et al., 1998; John, 1991; Redmond et al., 2004), and up to1.3 wt.% Cu and 0.3 wt.% Fe (Klemmet al., 2007; Rusk et al., 2004, 2008;Sawkins and Scherkenbach, 1981). These high observed metal solubil-ities are consistent with or exceed experimental observations andtheoretical predictions based on chloride complexing alone (e.g.,Candela and Holland, 1984, 1986), and suggest that other volatileligands such as sulfide speciesmay enhance the solubility of chalcophileelements such as Cu and Au in high temperature aqueous fluids (e.g.,Heinrich et al., 1992; Pokrovski et al., 2005, 2008; Seo et al., 2009; Simonet al., 2006; Zajacz et al., 2008, 2011).

Shallowly emplaced magmas will exsolve fluids under pressure–temperature (P–T) conditions that lie within the two-phase liquid–vapor field for the bulk fluid composition, resulting in immediateformation of an immiscible low salinity vapor and high salinity brine(Fig. 6a).

The critical point in the H2O–NaCl system (which is commonlyused as a proxy for magmatic fluids) occurs between ~1.0 and 1.4 kbfor fluid temperatures between 600° and 800 °C (the typicaltemperature range for fluids exsolved from intermediate to felsicmagmas) (Pitzer and Pabalan, 1986; Sourirajan and Kennedy, 1962),which is equivalent to depths of between ~3 and 5 km at lithostaticpressures. Most porphyry deposits and their host plutons areemplaced at depths between 1 and 6 km (Seedorff et al., 2005), sofluids exsolving directly from magmas at these depths will typicallyform immiscible liquid and vapor plumes (e.g., Henley and McNabb,1978; Nash, 1976). However, the bulk of the fluids and metals inporphyry deposits are likely initially sourced at deeper levels (5–10 km, as noted above) from larger volumes of magma in mid- toupper crustal batholithic complexes (Candela and Piccoli, 2005; Cloos,2001; Damon, 1986; John, 1991; Richards, 2003, 2005; Shinohara andHedenquist, 1997), at which depths the fluids will be supercritical. Asthese deep fluids rise into shallow cupola zones extending above themain batholith, they will likely intersect the solvus on the vapor side,and will begin to condense a dense saline brine (Ahmad and Rose,1980; Figs. 6a and 7).

Supercritical fluids are highlymobile (e.g., Coumou et al., 2008; Dunnand Hardee, 1981; Norton and Dutrow, 2001) and behave differently interms of magma–fluid partitioning compared to fluids exsolved atshallower depths in the two-phase field. In particular, Henley andMcNabb (1978) suggested that the higher density and viscosity of salinebrine condensates might restrict their flow, leaving them as a denseresidual liquid in the deeper parts of evolving magmatic hydrothermalsystems (see also: Lewis and Lowell, 2009;White et al., 1971). The lowerdensity vapor or supercritical fluid would be expected to be highlyupwardlymobile, aswell as larger in termsof bothvolumeandmass than

the brine phase (assuming an initial bulk salinity below ~20 wt.% NaCl),and sohasmuchgreater potential as anefficient transportingmediumforore components. However, until fairly recently, it was assumed that lowsalinity vaporswould not have the capacity to dissolve large quantities ofbasemetals as chloride complexes, and the brine phasewas therefore thefavored ore-forming medium (e.g., Bodnar and Beane, 1980; Cline andBodnar, 1991; Eastoe, 1982; Hedenquist and Richards, 1998; Moore andNash, 1974; Nash, 1976; Shinohara, 1994; Williams et al., 1995).Observations of chalcopyrite crystals trapped in some vapor-rich fluidinclusions (e.g., Bodnar and Beane, 1980) were explained by some asproducts of heterogeneous trapping (see discussion in Mavrogenes and

Fig. 7. Schematic cross-section through a typical coupled arc batholith–cupola–volcanic system, with associated porphyry Cu±Au and linked high sulfidation Cu–Au epithermaldeposits. Also shown are the thermal structure, fluid flow pathways and characteristics during the main stage of hydrothermal activity, and overlapping hydrothermal alterationzones. Propylitic alteration by circulating heated groundwaters can be assumed to affect all the supracrustal rocks in the field of view, with greatest intensity (epidote, actinolite)close to the intrusions, fading to background distally. Modified from Richards (2005); sources: Sillitoe (1973, 2010), Dilles (1987), Shinohara and Hedenquist (1997), Hedenquist etal. (1998), and Fournier (1999).

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Bodnar, 1994). The advent of quantitative analysis of single fluidinclusions by synchrotron X-ray microprobe, proton-induced X-rayemission spectroscopy (PIXE), and laser ablation-inductively coupledplasma-mass spectrometry (LA-ICP-MS) revealed that, contrary to theseearlier assumptions, considerable amounts ofmetal, including CuandAu,were indeed consistently present in some vapor-rich inclusions (Audétatet al., 2000; Harris et al., 2003; Heinrich et al., 1992, 1999; Klemm et al.,2007; Lowenstern et al., 1991; Simon et al., 2005, 2006; Ulrich et al.,2001), and subsequent work has shown that much of this metal contentis transported as sulfide complexes rather than as (or in addition to)chloride complexes (Cauzid et al., 2007; Heinrich et al., 1992, 1999;Pokrovski et al., 2005, 2008; Seoet al., 2009; Simonet al., 2006;Zajacz andHalter, 2009; Zajacz et al., 2010).

Early resistance to these ideas was, I believe, at least in part due toconfusions of terminology: most papers dealing with this subject havereferred tometal solubility and transport in a vapor phase, suggesting tothe unwary a lowdensity, low salinity gas,whereas for themost part thefluids in question were either single-phase supercritical fluids, orrelatively highdensity vapors just below their critical points. Under suchhigh P–T conditions, vapors are almost as saline as the initial single-phase fluid from which they evolved, and therefore they still containplenty of chloride for base metal complexing. But perhaps moreimportantly, as noted above, these vapors will also contain a highconcentration of volatile sulfur species, which now appear to beessential for the efficient solvation of chalcophile elements under highP–T conditions.When combinedwith the highermass proportion of thevapor phase (versus brine condensate) and its highmobility, amodel fortransport of the bulk of the metal flux in porphyries by relatively densevapors or supercritical fluids is now well established (Klemm et al.,2007; Landtwing et al., 2010; Williams-Jones and Heinrich, 2005).

The rapid reduction in the efficiency of transport of metals as thevapor plume rises, cools, and becomes less dense by brine condensation,

may explain precipitation of the bulk of Cu, Mo, and some Au over arelatively narrow temperature interval between 425°–320 °C (Hemley etal., 1992;Klemmet al., 2007; Landtwing et al., 2005). At shallower depthsand lower temperatures, the vapor phase may become too dilute totransport significant amounts of base metals as chloride complexes, butmay continue to carry some metals such as Au, Cu, As, and Sb as sulfidecomplexes (Deditius et al., 2009; Simon et al., 2006, 2007), eventuallyeither venting them to the surface in high-temperature fumaroles (e.g.,Chaplygin et al., 2007; Hedenquist et al., 1994a; Symonds et al., 1987;Tarana et al., 1995; Tessalina et al., 2008) or precipitating them in thenear-surfacehigh-sulfidationepithermal environment (see Section4.2.1;Deditius et al., 2009; Hedenquist et al., 1993, 1994b; Heinrich et al., 1999,2004; Larocque et al., 2008; Murakami et al., 2010; Pudack et al., 2009;Williams-Jones and Heinrich, 2005).

4. Magmatic–hydrothermal ore formation

The focus of this paper is on the flux of metals in subduction-related magmatic systems, but this would be of little practical interestif that flux did not ultimately lead to ore formation. Thus far, we havefocused on the importance of firstly not losing significant amounts ofmetal to a fractionating or residual sulfide phase, and then efficientlypartitioning those metals into a highly mobile aqueous fluid phase.What subsequently happens to that fluid phase dictates whethereconomic concentrations of metals are precipitated (grade), whereasthe scale of the magmatic and derivative hydrothermal systemcontrols the total amount of metals precipitated (tonnage).

4.1. Porphyry Cu ore formation

In a landmark paper, Cline and Bodnar (1991) presented a modelfor the evolution of magmatic-hydrothermal systems from initial

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aqueous fluid exsolution to cooling and mineral precipitation. A keyfinding of this work was that, although there are many variables thatcan affect the specific evolutionary path of a given system, aneconomic porphyry Cu deposit can potentially be formed from quitesmall volumes of typical andesitic arc magma. For example, theauthors showed that 15–90 km3 of andesitic magma initially contain-ing 50 ppm Cu, 2.5 wt.% H2O, and with a Cl/H2O ratio=1, is sufficientto generate a 250 Mt deposit with a grade of 0.75 wt.% Cu. The range ofrequired magma volumes reflects variables such as the depth ofmagma emplacement and the compatibility of Cu with fractionatingmineral phases (larger volumes are required if Cu is compatible withearly fractionating silicate, oxide, or sulfide minerals). Details of themodel, later updated in Cline (1995), show that magmas emplaced atmoderate depths (1.0 to 2.0 kb, equivalent to depths of ~4–8 km)most efficiently partition Cu into early saturating saline fluids,whereas Cu and Cl are only released from shallowly emplaced(0.5 kb, or ~2 km) magmas during late stages of crystallization (seealso Candela, 1989b). (Note that these depths reflect the locus of fluidand metal exsolution from the magma under lithostatic pressureconditions, as opposed to the depth of subsequent hydrothermalmetal deposition, which will be at shallower levels and likely underhydrostatic pressure conditions; Fournier, 1999.) In large measure,these differences reflect the change in properties of the magmatic–hydrothermal fluid, which will separate initially as a moderatelysaline supercritical fluid in deeper systems (Fig. 6a, path 1), but asimmiscible vapor and brine at shallower levels (Fig. 6a, path 2). Cline(1995) concluded that optimum conditions for porphyry Cu oreformation are obtained where magma containing ~4 wt.% H2O andwith a high initial Cl/H2O (typical of most arc magmas) is emplaced atmoderate crustal depths, thereby maximizing the efficiency of Cupartitioning into an early saline fluid phase.

The results of these studies indicate that no special conditions ormagmatic metal enrichment are required to form even large porphyryCu deposits. Instead, the question is rather one of process efficiency:where (what depth) and how are metalliferous fluids released andchanneled? This finding is of fundamental importance from anexploration perspective, because it shifts the focus from seekinganomalously metal-rich magmas (which have proven elusive; Jenneret al., 2010) to searching for optimal ore depositional settings inotherwise normal tectonomagmatic environments (e.g., Tosdal andRichards, 2001).

Cloos (2001), Shinohara et al. (1995), and Shinohara andHedenquist(1997) considered the question of process efficiency from theperspective of physical separation and focusing of volatile releasefrom the magma. Cloos (2001) suggested that the classic cupola shape(Norton, 1982) of porphyry systems above larger batholithic magmachambers reflects convective circulation of bubbly magma into theshallowapical parts of these systems (at 1–3 kmdepth), where volatilesphysically separate from the melt and coalesce to form a discrete fluid-filled cupola. The now-dense, volatile-depleted magma forms adownward return flow to complete the convective cycle. In contrast,Shinohara and Hedenquist (1997) envisaged fluids physically separat-ing from the magma at greater depth within the underlying magmachamber, and rising as a discrete plume up apical channelways formedby fractures and dikes in the brittle carapace (Fig. 7).

A key aspect of both of thesemodels is that volatile separation, andCu partitioning, occurs from a much larger volume of magma(emplaced at deeper levels) than that preserved and commonlyvisible within the shallow-level ore body. In Cloos's (2001) model,volatile-rich, bubbly magma rising from the underlying batholithconvects through the cupola zone where it releases its fluids, whereasin Shinohara and Hedenquist's (1997) model, vesiculation andconvective circulation occur in the underlying magma chamber itself,and fluids are released as a plume into the base of the apical dikesystem. Combining these models with (Cline, 1995; Cline andBodnar's, 1991) calculations suggests that maximum ore-forming

efficiency in porphyry systems is likely achieved where volatilesaturation occurs in large (≥100 km3) mid- to upper crustal magmachambers at depths ≥6 km, containing moderately hydrous (N4 wt.%H2O) and Cl-rich magmas. These fluids either rise as bubbly magma oras a separate volatile plume into the apical parts of the systemwhere decreasing pressure and temperature cause deposition of Cuand Mo±Au (Candela, 1989b; Shinohara et al., 1995).

Focusing of magma ascent and fluid flow into narrow apicalregions, or cupolas, is likely to be a function of structure in the brittlerocks overlying the batholithic system (Tosdal and Richards, 2001,and references therein). Shallow crustal magma emplacement willcause extensional doming in the cover rocks, with dilational faultzones providing high-permeability pathways for fluid and magmaascent (Burnham, 1979). Evidence from the dike emplacementliterature suggests that such fractures may first be opened andpropagated by volatile pressure, and only later filled by more viscousmagma (Burnham, 1979; Carrigan et al., 1992; Rubin, 1995). Thisraises the intriguing possibility that the cylindrical shapes of manyporphyry stocks may have arisen first as breccia pipes or diatremesbored out by rapidly escaping volatiles, only later to be back-filledwith porphyritic magma (Fig. 8; Norton and Cathles, 1973; see alsoFig. 2 in Anderson et al., 2009, and Fig. 8 in Sillitoe, 2010, and Fig. 17 inVry et al., 2010).

To some extent, focusing of magma and fluid flow along narrowconduits may be self-organizational, because once initial channelways have developed, they will represent high-permeability path-ways and are likely to thermally weaken the wall rocks and promotefurther fracturing and channeling.

Narrow focusing of apical fracturing and subsequent fluid flow arelikely to be critical to the formation of high grade porphyry deposits(e.g., El Teniente, Chile; Vry et al., 2010), while multiple breccia/intrusive events can potentially increase tonnage (provided that laterevents do not destroy earlier mineralization).

Given suitable fluid flow focusing, three further factors combine tocause maximum efficiency of metal deposition within relatively smallvolumes in porphyry Cu±Mo±Au deposits. All three effects arerelated to the steep temperature gradient in the cupola zone (Fig. 7),and they therefore control the vertical range of ore deposition. On theother hand, fluid focusing controls the lateral extent of mineralization.In combination, the highest grades will occur where ore deposition isboth focused laterally and restricted vertically.

The first factor is that Cu solubility (as chloride species) decreasesdramatically as fluids cool through the temperature interval ~400° to300 °C (Crerar and Barnes, 1976; Hemley et al., 1992; Klemm et al.,2007; Landtwing et al., 2005; Xiao et al., 1998). Given the very sharptemperature gradient implied by near-surface emplacement ofmagma (Fig. 7), this temperature interval will correspond to a narrowdepth range, likely with 1 or 2 km of the surface (although it mayextend to greater depths with time as the magmatic–hydrothermalsystem begins to cool). This depth range is typical for ore formation inmany porphyry systems.

The second factor is that SO2 dissolved in the magmatic–hydrothermal fluid phase progressively disproportionates to H2Sand H2SO4 as the fluid cools below ~400 °C (Holland, 1965; Kusakabeet al., 2000; Reeves et al., 2010; Sakai and Matsubaya, 1977):

4SO2 + 4H2O ⇔ H2S + 3H2SO4: ð1Þ

This reaction generates both hydrogen sulfide, which initiatesabundant precipitation of sulfide minerals (i.e., chalcopyrite, pyrite,molybdenite), and also sulfuric acid, which causes early deposition oflarge volumes of anhydrite in the potassic alteration zone, andprogressively increasing degrees of hydrolytic alteration (an initialshift from feldspar-stable potassic alteration, to muscovite/sericite-stable phyllic alteration). Consequently, the bulk of Cu-sulfide miner-alization occurs at the low-temperature, late-stage end of the potassic

Porphyryintrusion

Mixed magmatic-hydrothermal

breccia

Cuspateintrusiveborder

Porphyryintrudingcomagmaticbreccia

Cuspateporphyry

clastMixed magmatic-

hydrothermalbreccia

Hydrothermalquartz fillingcavity

Limestonewallrock

clast

Porphyrydike

(a) (b)

(c)

~20 cm

Fig. 8. Photographs of porphyritic magma invading co-magmatic hydrothermal breccia pipe (a, c) and breccia vein/dike (b) from the Pachapaqui Ag–Cu–Pb–Zn deposit, Péru. Thebreccia consists of fragments of porphyry magma and country rock; note scalloped, cuspate margins of the porphyry body in (a, b) and porphyry clasts in (c), indicating that themagma was molten at the time of breccia formation and subsequent intrusion into the breccias. In (b), large vuggy spaces are partially filled with hydrothermal quartz, reflecting therole of fluids in breccia formation. Scale in (c) is in centimeters.

15J.P. Richards / Ore Geology Reviews 40 (2011) 1–26

alteration phase, just prior to the onset of phyllic alteration (i.e.,corresponding to the “ore shell” of the classic Lowell and Guilbert, 1970,porphyry Cu model).

The third factor is the increased permeability over this tempera-ture range caused by a combination of the transition in silicate rocksfrom ductile to brittle behavior at temperatures (between ~400°–350 °C; i.e., the brittle–ductile transition; Fig. 7; Cathles, 1991;Fournier, 1999; Landtwing et al., 2005), and a window of retrogradesilica solubility (between ~550–350 °C; Fournier, 1985). Not only dothese processes result in the formation of open-standing brittle veinsand porosity, thereby facilitating rapid upward fluid flow andwallrock permeation (e.g., the crackle breccias, stockworks, anddisseminated mineralization textures so characteristic of porphyry Cudeposits), but this transition also represents the boundary betweenlithostatic and hydrostatic fluid pressures (a pressure differential of~3×). Sudden depressurization of fluids across this boundary can beexpected to have major effects on fluid properties, including phaseseparation (Fig. 6) and consequent changes in metal solubility (e.g.,Landtwing et al., 2005, 2010; Murakami et al., 2010).

In combination, these four factors, (1) spatial focusingoffluidflow innarrow cupolas, (2) reduction of metal solubility, (3) increaseddissolved sulfideactivity, and (4)permeability increasedue to transitionfrom ductile to brittle fracturing and retrograde silica solubility (with a

large pressure drop), serve to narrowly focus Cu-sulfide mineralizationboth laterally and vertically within cupola zones above large mid- toupper crustal batholithic complexes. The most likely reasons forotherwise prospective porphyry systems to be unproductive will beeither a failure to focus fluid flow, or simply insufficient fluid supply(likely due to an insufficiently large underlying magmatic system).

4.2. Epithermal Cu–Au ore formation

4.2.1. High-sulfidation epithermal Cu–Au depositsAs noted in Section 3.1, although the bulk of Cu (andMo) appears to

beprecipitated over the temperature interval 425°–320 °C, someCu andother metals such as Au, Sb, and As may remain in solution as sulfidecomplexes, to be carried into the shallow epithermal regime. Observa-tions from rare fluid inclusions in high-sulfidation epithermal Cu–Audeposits suggest that the ore-forming fluid was a low- to moderate-salinity liquid (0.2 to 4.5 wt.%NaCl equivalent; Hedenquist et al., 1994b;Mancano and Campbell, 1995), which paragenetically post-datesadvanced argillic alteration formed by highly acidic magmatic gasses(Arribas, 1995; Hedenquist et al., 1994b; Stoffregen, 1987). Thisapparent inconsistency has been explained by Hedenquist et al.(1998), Heinrich et al. (2004), and Heinrich (2005) in terms ofcontraction of a moderate salinity supercritical magmatic fluid or

16 J.P. Richards / Ore Geology Reviews 40 (2011) 1–26

vapor by rapid cooling during ascent such that it does not touch, orbarely touches, the two-phase solvus (Figs. 6b and 9, paths 3 and 4,respectively). Because of the curvature of the solvus crest (critical curve)to lower salinities at low temperatures and pressures, this moderatelysaline fluid will lie on the liquid side of the solvus at shallow depths,although it may have contracted from what was originally a vapor orsupercritical fluid phase at higher temperatures and depths.

Heinrich et al. (2004) specifically suggested that brief intersectionwith the solvus at high pressures and temperatures (Figs. 6b and 9, path4)might be an importantway to shed chloride-complexed componentssuchas Fe fromthe vapor in a brine condensate, leavingbisulfide ligandsfree to bond with Cu and Au in the residual vapor and transport it toshallower (epithermal) levels. Thismodel requires that the vapor leavesthe two-phase surface again by coolingat pressure, therebypassingoverthe crest of the solvus (the critical curve; (Figs. 6b and 9, path 4); uponsubsequent ascent and depressurization, this fluid will be a liquid, asdescribed above. Heinrich et al. (2004) argued that the Fe-rich brinecondensation step is essential for retention of Au (and Cu) in the vaporphase, because otherwise Fe will tend to precipitate as Cu–Fe-sulfideminerals and strip the fluid of bisulfide ligands, thus causing Au to co-precipitate at depth (a possible mechanism for the formation ofporphyry Cu–Au deposits). However, this condensation step mustthen be followed by cooling while still at depth, in order to “lift” thevapor phaseoff the two-phase surface, and allow it to contract to a liquid(Figs. 6b and 9, path 4).

Epithermal

V-L solvus (vapour)

V-L solvus (liquid)

Brit

tle(h

ydro

stat

ic)

Ductile

(lithostatic)

30°C

/km

geo

ther

m

V

Main-stagebrittle

porphyryveins

T (°C)0 100 200 300 400 500

P (

bars

)

500

0

1000

1500

2000

0 100 200 300 400 500

Fig. 9. Pressure (depth)–temperature section through the H2O–NaCl phase diagram, with vaDriesner and Heinrich, 2007). Red curves indicate that the solvus phase is a vapor, blue ccomposition. Also shown are the wet granite and granodiorite solidi (Burnham, 1979), andgradients (Barbier, 2002; Goff et al., 1992; Noorollahi et al., 2007). Depths are indicated forranges for supercritical magmatic fluid exsolution, early high-temperature porphyry veins, lafluid P–T paths are shown corresponding to: (1) a typical porphyry-forming fluid path; (2contraction path of Hedenquist et al. (1998); and (4) the contraction with minor brine con

A more detailed analysis of the phase diagram shown in Fig. 6reveals that a typical magmatic fluid of 2–13 wt.% NaCl would have tofirst intersect the solvus at temperatures above ~400°–500 °C for thismodel to work optimally, otherwise it would fall on the liquid side ofthe two-phase field, and would become more saline by boiling off adilute vapor (Figs. 9 and 10). Assuming that this did indeed happen,then the smaller (by mass) vapor component would have to leave thetwo-phase surface again at pressure before cooling below ~400 °C,otherwise its salinity rapidly falls to sub-weight percent levels atlower temperatures and lower pressures (Figs. 6 and 10), which areinconsistent with fluid inclusion evidence for low tomoderately salineliquids in high-sulfidation epithermal ore formation (Hedenquist etal., 1994b;Mancano and Campbell, 1995). Thus, the P–T trajectory of afluid that satisfies Heinrich et al.'s (2004) model for high-sulfidationepithermal Cu–Aumineralization (Figs. 6b and 9, path 4) is somewhatunique, and may occur only rarely or fleetingly during the waningstages of a cooling magmatic–hydrothermal system.

A more normal, or early ascent pathway for a magmatic–hydrothermal fluid would be for it to rise more-or-less is enthalpicallyor quasi-adiabatically along a steep P–T gradient (e.g., Hemley andHunt, 1992; Henley and Hughes, 2000; Wood and Spera, 1984), andtherefore to dive deeply into the two-phase field and separate into anincreasingly dilute vapor phase and a saline brine (Figs. 6 and 9, path1), or even to boil dry to halite plus vapor (Figs. 6 and 9, path 2). Suchfluid pathways might be consistent with the widespread and intense

sudil

oseti

narg

teW

sudil

oseti

roid

onar

gte

W

Earlyporphyry

veins(2-phase

fluid)

Deepmagmatic

fluidexsolution

3

41 1 wt.%

5 wt.%

10 w

t.%

20 w

t.%

300°C/km

geotherm

+Halite

V/L+Halite

Shallowmagmatic

fluidexsolution

2

Deepsingle-phasefluid

Depth (km

)

1 (hyd)

2 (hyd)1 (lith)

2 (lith)

8 (hyd) 3 (lith)

4 (lith)

5 (lith)

6 (lith)

7 (lith)

8 (lith)

3 (hyd)

4 (hyd)

5 (hyd)

6 (hyd)

7 (hyd)

9 (hyd)

10 (hyd)

0600 700 800 900 1000

600 700 800 900 1000

11 (hyd)

12 (hyd)

13 (hyd)

14 (hyd)

15 (hyd)

pour–liquid (V–L) solvi drawn for 1, 5, 10, and 20 wt.% NaCl (data from Driesner, 2007;urves that it is a liquid; the transition point corresponds to the critical point for thataverage crustal (30 °C/km) and high (300 °C/km, near active volcanism) geothermal

hydrostatic (hyd) and lithostatic (lith) pressure conditions. Typical temperature–depthter main-stage brittle porphyry veins, and epithermal mineralization are indicated. Four) a shallow high-temperature path boiling to dryness (V+Halite field); (3) the deepdensation path of Heinrich et al. (2004).

Crit

ical

cur

ve

L+H

(40

0°C

)V+H (400°C)

V+L(700°C)

V

wt.% NaCllog (wt.% NaCl)

P (

bars

)

scalechange

100

0

200

300

400

500

600

700

800

900

1000

1100

1200

1300700°C

600°C

500°C

400°C

Singlephasemag-maticfluid

V+L(600°C)

V+L(400°C)

V+L(500°C)

V+H (700°C)

V+H (600°C)

V+H (500°C)

L+H

(50

0°C

)

L+H

(70

0°C

)L+H

(60

0°C

)

Critical Pof H2O L

(400

°C)

L (5

00°C

)

L (7

00°C

)

L (6

00°C

)-8 -7 -6 -5 -4 -3 -2 -1 0/1 10 20 30 40 50 60 70 80 90 100

100

0

200

300

400

500

600

700

800

900

1000

1100

1200

1300-8 -7 -6 -5 -4 -3 -2 -1 0 10 20 30 40 50 60 70 80 90 100

V-L solvus (liquid)

V-L solvus (vapour)

V-H solvus (vapour)

Fig. 10. Pressure–XNaCl section through the H2O–NaCl phase diagram, with vapor–liquid (V–L) solvi drawn at various temperatures (data from Driesner, 2007; Driesner and Heinrich,2007). Red curves indicate that the solvus phase is a vapor, blue curves that it is a liquid; below the vapor–halite (V–H) solvus, vapor curves are shown in orange. Note the scalechange on the salinity axis at 1 wt.% NaCl, in order to illustrate the extremely low salinity of low-temperature vapor phases. The range of salinity for typical deeply exsolved single-phase (supercritical) magmatic fluids (2–13 wt.% NaCl) is shown in gray. Fluids that intersect the V–L solvus above ~400°–500 °C will be moderate-density vapors and will condensea small amount of dense, saline liquid; fluids that intersect the solvus below this temperature will be liquids and will boil off a dilute, low-density vapor phase.

17J.P. Richards / Ore Geology Reviews 40 (2011) 1–26

acidic (advanced argillic) alteration commonly found at shallow levelsabove porphyry systems, which is caused by acidic gasses (H2SO4,HCl) condensing from a low-density vapor plume. As Heinrich et al.(2004) noted, such acidic vapors do not appear to transport Au (or Cu)effectively because bisulfide (HS–) ligands are hydrolized to H2S. Thismay explain why advanced argillic alteration caps are commonlybarren (in terms of Au and Cu) and merely generate permeability thatpotentially focuses later ore-forming fluid flow. Thus, mineralizationmay only occur where later moderate salinity liquids have followed ahigher pressure, rather specialized cooling path, as described above(Heinrich et al., 2004).

4.2.2. Low-sulfidation epithermal Au deposits (including alkalic-typedeposits)

Although low-sulfidation epithermal Au deposits are commonlyfound in volcanic terrains, their link to magmatism is more tenuousthan high-sulfidation deposits, and there is commonly evidence for agreater involvement of meteoric groundwater in their formation thanmagmatic fluids (e.g., Faure et al., 2002; Field and Fifarek, 1985; Healdet al., 1987). Nevertheless, alkalic-type epithermal Au deposits, whichare mineralogically similar to low-sulfidation deposits (adularia andsericite — or roscoelite [vanadium mica] — are stable), do show astrong temporal and genetic relationship to alkalic magmas, typicallyin relatively small and isolated intrusive complexes located in back-arc or post-subduction settings (e.g., Jensen and Barton, 2000; Kelley

et al., 1998; Müller and Groves, 1993; Mutschler et al., 1985; Richards,1995; Thompson et al., 1985).

Fluids in these alkalic-type deposits are commonly low- tomoderate salinity (0–10 wt.% NaCl equiv.), low temperature (typically≤250 °C) liquids, with evidence for decompressional boiling or fluidmixing as the prime ore depositional mechanism in high gradebreccias and veins (Jensen and Barton, 2000; Richards, 1995). Stableisotopic compositions of these fluids are generally ambiguous, andpermit interpretations of the involvement of either isotopicallyexchanged meteoric waters or magmatic fluids, or both (Ahmad etal., 1987; Carman, 2003; Richards, 1995; Richards and Kerrich, 1993;Ronacher et al., 2004; Scherbarth and Spry, 2006; Zhang and Spry,1994). Similar stable isotopic data from other low-sulfidation depositsare commonly interpreted to reflect a meteoric fluid source because ofthe absence of clearly coeval magmatism (as noted above). However,the close link with magmatism in alkalic-type systems suggests that amagmatic fluid source is more likely (e.g., Carman, 2003; Scherbarthand Spry, 2006; Simmons and Brown, 2007), and the vaporcontraction model described in Section 4.2.1 could explain theobserved characteristics of these ore fluids (i.e., direct contraction toa moderate salinity liquid from a high temperature magmatic fluid atdepth). Because of the association of these deposits with relativelysmall intrusive complexes, and therefore a smaller crustal thermalanomaly, a shallower fluid P–T path with cooling at depth is morelikely, consistent with a model of vapor contraction.

18 J.P. Richards / Ore Geology Reviews 40 (2011) 1–26

5. Summary and conclusions

5.1. Sources of magmas and metals

Magmatic–hydrothermal porphyry Cu±Mo±Au, Au,Mo, and Sn–Wdeposits (and related epithermal Au deposits), derive their metals fromtheir associatedmagmas.With the exception of porphyry Sn–Wdepositsthat are associated with crustally derived S-type granites, mostother deposits in this grouping are formed by calc-alkaline to mildlyalkaline I-type granitoids directly or indirectly related to subduction.Sources of these magmas include subduction-metasomatized astheno-spheric mantle wedge, basaltic oceanic crust and/or seafloor sediments,and, in post-subduction settings, subduction-modified upper platelithosphere. The majority of normal arc porphyry systems are generatedfrom hydrous mantle wedge melts that have interacted to varyingdegrees with the upper plate lithosphere during passage towards thesurface. Assimilation and fractional crystallization processes fundamen-tally change the composition of the primary basaltic arc magmas tointermediate calc-alkaline compositions, with fractionated trace elementpatterns and evolved (crustal) isotopic signatures.

Seafloor sediments and basaltic oceanic crust only melt underunusually hot subduction zone conditions or locally at plate edges, anddespite widespread claims of the identification of slab melts (adakites)in the literature based on high Sr/Y and La/Yb ratios in some evolvedgranitoids, this process is unlikely to be a major contributor to arcmagmatism and metallogeny. Rather, these subtle trace element ratioscan be readily explained by fractionation of hornblende±titanite andresidual garnet, and suppression of plagioclase fractionation fromwater-rich mantle wedge basaltic magmas. These trace elementcharacteristics are therefore an indicator of high magmatic watercontent rather than being a source signature, and this likely explains thecommon association of high-Sr/Y (i.e., hydrous) magmas with mag-matic–hydrothermal ore deposits — without magmatic water, suchdeposits cannot form.

Potential sources of metals in arc magmas include the oceanic crustand sediments (via dehydration fluids or melting), the mantle wedge,and theupper plate crust. Fluid–mobile elements suchasK,Rb, Cs, Ca, Sr,Ba, U, B, Pb, As, Sb, Tl, and possibly Cu, Au, Re, and the Pd-groupelements, alongwith large amounts ofH2O, Cl, and S, arefluxed from thedehydrating subducting slab into the mantle wedge, causing metaso-matism and partialmelting by lowering the peridotite solidus. Althoughthere is some evidence for slab-derived fluid contributions of chalco-phile and highly siderophile elements (Cu, Au, PGE) to the mantlewedge, it is not clear that this is a necessarymetallogenic step, the uppermantle already containing significant amounts of these elements. Likely,a more important control on the metal content of subsequent partialmelts is the abundance and stability of residual sulfide phases in theasthenospheric mantle source. Under the high fO2 and fS2 conditions ofarcmagmatism, sulfur will be dominantly present as sulfate and sulfate,but saturation in small amounts of sulfide phases is also likely. Thesesulfide phases will tend to deplete the magma in highly siderophileelements (Au and PGE), but will not be present in sufficient volume tosignificantly deplete themagma inmore abundant chalcophile elementssuch as Cu and Mo. Such magmas therefore have the potentialsubsequently to form porphyry Cu–Mo deposits. Thus, Cu (and perhapsMo) are thought to be predominantly derived from the mantle, plus orminus contributions from the subducting slab. Gold and silverpresent inminor amounts in such deposits may also be derived from subductionsources, although there is some evidence for additional contributionsfrom the upper plate crust (especially for Ag, and also perhaps Mo).

Arc-like magmas and related porphyry and epithermal ore depositsalso occur in post-subduction tectonic settings, such as subductionreversal or migration, arc collision, continent–continent collision, andpost-collisional rifting. They are distinguished from normal subduction-related suites by slightly higher magmatic alkali (K2O and Na2O)contents, and by the occurrence of Au-rich deposits (although normal

porphyryCu–Modeposits can alsooccur). Suchmagmatic–metallogenicsystems are thought to form by remelting of previously subduction-modified upper plate lithosphere, and in particular the lower crustalamphibolitic cumulate roots of former arc magmatic complexes.Remelting can be triggered by crustal thickening and thermal reboundfollowing arc or continent collision, delamination of sub-continentalmantle lithosphere causing direct exposure of the lower crust toasthenospheric temperatures and melts, and asthenospheric upwellingduring rifting of former arc crust. Sparse sulfide phases in these arccumulates, residual from fractionation of previous arc magmas, willlikely be rich in chalcophile and highly siderophile elements. Duringlow-volumemelting under relatively low fS2 conditions (in the absenceof a flux of S from active subduction), these sulfide phases will likelyredissolve in the mildly alkaline partial melt, and may provide a sourcefor Au-rich (±PGE) post-subduction porphyry Cu–Au and epithermalAu systems: examples include the Roşia Montană, Skouries, Kisladag,Çöpler, and Sari Gunay porphyry Cu–Au and epithermal Au deposits inthe Neo-Tethyan belt of Romania, Greece, Turkey, and Iran, and theGrasberg, Ok Tedi, Porgera, Lihir, and Emperor porphyry Cu–Au andepithermal Au deposits in the southwest Pacific. However, goldenrichments may not occur where more abundant sulfides werepresent in the former arc complex, leading to more “normal” porphyryCu±Mo systems: examples include the Kerman porphyry Cu belt ofcentral Iran, and the Gangdese porphyry Cu belt of Tibet.

Porphyry Mo and Sn–W deposits associated with felsic magmas incontinental interiors are thought to form mainly by partial melting ofcontinental crust during rifting to form S-type, lithophile element-richgranitic magmas. A role for mafic, mantle-derivedmagmas is suggestedby the common association with such rocks, but their role may bepredominantly as a heat source for crustal melting and a trigger forvolatile saturation and eruption, rather than as a unique source ofmetals.

5.2. Porphyry and epithermal ore formation

In the porphyry and epithermal ore depositional environment, acritical role is played by aqueous fluids exsolving from hydrousmagmas emplaced in themid- to upper crust. The P–T–X properties ofmagmatic hydrothermal fluids, approximated by the H2O–NaClsystem, combinedwith volatile solubility in intermediate compositionmagmas, suggest that fluid saturation occurs at depths of 5–10 km inthe batholithic roots of arc magmatic systems. Volatile exsolutionmaylead to the upward propagation of buoyant bubbly magma as dikesand stocks intruded into the overlying shallow crust (with or withoutsubsequent eruption at surface), and/or the rapid ascent of a separatevolatile plume. Structural focusing and chanelling of these evolvedmagmas and fluids creates a cupola zone characterized by highthermal gradients and fluid flux. Metals (Cu, Mo, Au), which partitionstrongly into the saline (2–13 wt.% NaCl equivalent) and S-richmagmatic hydrothermal phase at high P and T in the underlyingbatholithic magma chamber, experience rapid reduction in solubilityas these fluids ascend, depressurize, and cool, with the bulk of Cu andMo being precipitated over a temperature range of 425°–320 °C at 1–6 km (commonly ≤2 km) depth.

This depth–temperature interval is critical because it also repre-sents: (1) the upward transition from ductile to brittle behavior in thecover rocks (~400°–350 °C), which facilitates fracturing and rapid fluiddepressurization; (2) a window of retrograde silica solubility (~550°–350 °C), which enhances permeability and porosity for ore deposition;and (3) the temperature range (b400 °C) over which SO2 in the fluidphase begins to disproportionate to H2S and H2SO4, which causesprecipitation of sulfideminerals (chalcopyrite, pyrite,molybdenite, rarebornite). This reaction also generates increasingly acidic fluids, leadingto the characteristic progression from feldspar-stable alteration assem-blages (potassic), through muscovite/sericite-stable assemblages

19J.P. Richards / Ore Geology Reviews 40 (2011) 1–26

(phyllic), to clay- (argillic) and alunite-stable assemblages (advancedargillic).

Depending on the depth of exsolution, the initialmagmatic fluidwilleither be a supercriticalfluid (below~6 kmdepth) orwill exist as a two-phase moderate salinity vapor and high density brine. As the plumeascends, it will intersect its solvus (in the case of a deeply exsolvedsupercritical fluid), and the vapor phase will become progressively lesssaline through brine condensation. At the level of porphyry oreformation, both the brine and vapor phase may contribute to metaltransport and deposition, although there is increasing evidence for theimportance of the vapor phase as a large-volume, highly upwardlymobile transportationmedium. However, as this vapor phase continuesto ascend, it will rapidly decrease in salinity, such that chloride-complexed metals are unlikely to be transported by such fluids toshallow epithermal levels. It will also increase in acidity, therebyreducing the solubility of bisulfide-complexed metals such as Au (andperhaps also Cu) by protonation of HS– to H2S. This acidic vapor phase isresponsible for the extreme acid leaching in advanced argillic lithocapsabove porphyry systems.

High-sulfidation epithermal Cu–Au deposits are hosted by thesehighly permeable advanced argillic alteration zones, but appear to havebeen formed by later, moderately saline (0.2 to 4.5 wt.% NaClequivalent), less acidic liquids. Two models have been proposed toexplain the origin of these paragenetically late mineralizing fluids interms of contraction of a single phase (supercritical) magmatic fluid,with (Heinrich et al., 2004) or without (Hedenquist et al., 1998) briefintersection of the solvus. Because of the topology of the P–T–XNaCl

phase diagram, such fluids contract to a liquid phase upon cooling atpressure. Thus, these authors propose that high-sulfidation epithermalCu–Au deposits may be formed directly from late-stage magmatichydrothermal fluids.

Although low sulfidation epithermal Au deposits have not beendiscussed in detail in this paper, because most such deposits are notdirectly related to porphyry-type magmatic–hydrothermal systems,the vapor contraction mechanism might have applicability to alkalic-type epithermal gold deposits, which do show a close geneticrelationship to post-subduction alkalic magmas.

Acknowledgments

I would like to thank Nigel Cook and Timothy Horscroft for invitingme to submit this paper. Reviewpapers, by their nature, drawheavily onthe work of others, and I would particularly like to acknowledge thefollowing people who have influenced my thinking on this subject: P.Candela, J. Cline, J. Dilles, J. Hedenquist, C. Heinrich, R. Sillitoe, and R.Tosdal. An anonymous reviewer is thanked for helpful and constructivecomments. Thisworkwas funded by aDiscovery Grant from theNaturalSciences and Engineering Research Council of Canada.

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