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Metallogenic Provinces in an Evolving Geodynamic Framework

ROBERT KERRICH,†

Department of Geological Sciences, University of Saskatchewan, 114 Science Place, Saskatoon, Saskatchewan, Canada S7N 5E2

RICHARD J. GOLDFARB,U.S. Geological Survey, Box 25046, MS 964, Denver Federal Center, Denver, Colorado 80225-0046, and

Department of Geological Sciences, University of Colorado, 2200 Colorado Ave., Campus Box 399, Boulder, Colorado 80309

AND JEREMY P. RICHARDS

Department of Earth and Atmospheric Sciences, University of Alberta, Edmonton, Alberta, Canada T6G 2E3

AbstractThermal decay of Earth resulted in decreased mantle-plume intensity and temperature and consequently a

gradual reduction of abundant komatiitic basalt ocean plateaus at ~2.6 Ga. In the Neoarchean, ocean crust was~11 km thick at spreading centers, and abundant bimodal arc basalt-dacite magmatic edifices were constructedat convergent margins. Neoarchean greenstone belt orogenesis stemmed from multiple terrane accretion inCordilleran-style external orogens with multiple sutures, where oceanic plateaus captured arcs by jammingsubduction zones, and plateau crust melted to generate high thorium tonalite-trondhjemite-granodiorite suites.Archean cratons have a distinctive ~250- to 350-km-thick continental lithospheric mantle keel with buoyant re-fractory properties, resulting from coupling of the buoyant residue of deep plume melting to imbricatedplateau-arc crust. In contrast, Proterozoic and younger continental lithospheric mantle is <150 km thick,denser, and less refractory and therefore easily reworked in younger orogens. The supercontinent cycle has op-erated since ~2.8 Ga: Kenorland assembled at ~2.7 Ga, Columbia ~1.8 Ga, Rodinia ~1 Ga, and Pangea ~0.3Ga. Dispersal may have been triggered by superplumes.

Komatiite-hosted Ni deposits are related to plumes, where sulfide saturation resulted from crustal contam-ination. Base metal-rich volcanic rock-associated massive sulfide (VMS) deposits accumulated on thinned, frac-tured lithosphere within extensional oceanic suprasubduction environments, or back arcs, which were intrudedby anomalously hot subvolcanic sills; hence, their abundance in the Superior province of Canada (thick conti-nental lithosphere), contrasting with few in the Yilgarn craton of Australia (thick lithosphere). Orogenic golddeposits formed in sutures between accreted terranes associated with assembly of Kenorland. Diamonds werecreated by reaction of carbonate-rich asthenospheric liquids with continental lithospheric mantle at >240-kmdepth, mostly pre-2.7 Ga. They were entrained in kimberlitic to lamproitic melts related to superplume eventsat 480, 280, and ~100 Ma. Preservation of resulting mineral provinces stems from their location on stableArchean continental lithospheric mantle.

Decreased plume activity after 2.6 Ga caused sea level to fall, leading to the first extensive passive-marginsequences, including deposition of phosphorites, iron formations, and hydrocarbons, during dispersal ofKenorland from 2.4 to 2.2 Ga. Deposits of Cr-Ni-Cu-PGE were generated where plumes impinged on failedrifts at the transition from thick Archean to thinner Proterozoic continental lithospheric mantle, e.g., the GreatDyke, Zimbabwe, and later at Norilsk, Russia. Paleoproterozoic orogenic belts, for example, the Trans-Hudsonorogen in North America and the Barramundi orogen in Australia, welded together the new continent of Co-lumbia. Foreland basins associated with these orogens, containing reductants (graphitic schists) in the base-ment, led to the formation of unconformity U deposits, with multiple stages of mineralization generated fromdiagenetic brines for as much as 600 m.y. after sedimentation. Plume dispersal of Columbia at 1.6 to 1.4 Ga ledto SEDEX Pb-Zn deposits in intracontinental rifts of North America and Australia, extensive belts of RapakiviA-type granites on all continents, with associated Sn veins, and Fe oxide-Cu-Au-REE deposits. All were con-trolled by rifts at the transition from thick to thin continental lithospheric mantle. Plume impingement on Ro-dinia at ~1 Ga formed extensive belts of anorogenic anorthosites and Rapakivi granites in Laurentia andBaltica, the former hosting Fe-Ti-V deposits. Sedimentary rock-hosted Cu deposits formed in intracontinentalbasins from plume dispersal of Rodinia at ~800 Ma.

Iron formations and mantle plumes have common time series: Algoman type occur from 3.8 Ga to 40 Ma,granular iron formations precipitated on the passive margins of Kenorland at ~2.4 Ga, Superior-type formedon the passive margins of Laurentia, and Rapitan iron formations were created in rifts during latter stages ofdispersal of Rodinia at ~700 Ma. Accordingly, such deposits are not proxies for the activity of atmospheric O2.Rich Tertiary placer deposits of Ti-Zr-Hf, located on the passive margins of Australia and Southern Africa, re-flect multiple cannibalistic cycles from orogens that welded Rodinia and Pangea.

Orogenic Au deposits formed during Cordilleran-type orogens characterized by clockwise pressure-temper-ature-time paths from ~2.7 Ga to the Tertiary; Au-As-W and Hg-Sb deposits reflect the same ore fluids at pro-gressively shallower levels of terrane sutures. The MVT-type Pb-Zn deposits formed in foreland basins, with

† Corresponding author: e-mail, Robert.kerrich@usask.ca

©2005 Society of Economic Geologists, Inc.Economic Geology 100th Anniversary Volumepp. 1097–1136

Historical Perspective and Scope

Lindgren (1933) pioneered the concepts of both metallo-genic provinces and epochs. In the Economic Geology Fifti-eth Anniversary Volume, Turneaure (1955) synthesized globalmetallogenic provinces. He emphasized different classes ofore deposits, stable versus orogenic settings, lithologic ormagmatic associations of specific metal groupings, and therole of young mountain belts in preservation potential. Met-allogenic provinces of different ages were recognized, albeitwith large age uncertainties. Primary depositional setting ver-sus replacement was, and remains, an issue. Independently,Bilibin (1968) and Smirnov (1976) documented specific litho-tectonic and age associations for various classes of metallicdeposits in the former Soviet Union. Other comparative stud-ies of major ore provinces recognized the evolving crust-man-tle system as a control on lithological associations, magmaticstyle, and types of ore deposits (Pereira and Dixon, 1965;Stanton, 1972; Hutchinson, 1981). Atlases of the distributionof metallic deposits by geologic terrane and age were com-piled by Dixon (1979) and Derry (1980).

Meyer (1981) generated a global database of representativeor type metallic mineral deposits, and their age-lithotectonicassociation, in the Economic Geology Seventy-Fifth Anniver-sary Volume. He formulated the space-time distribution ofmetallogenic provinces in terms of two parameters: intervalsof geologic history during which specific classes of metallicdeposits formed, and changes of characteristics within a givenclass over the interval when that class formed. Meyer ob-served that trends of crustal evolution were not contempora-neous globally but did not cast his reviews in a plate tectoniccontext (Meyer, 1981, 1988).

The theory of plate tectonics was established in the 1970s,supplanting the geosynclinal concept of lithotectonic associa-tions (Kay, 1951; see Sengor, 1990, for a review). Elements ofthe theory included: recognition of ocean-floor spreadingfrom ages of volcanic islands and transform faults (Wilson,1965; Hess, 1968) and magnetic domains (Vine andMatthews, 1963), relative to mid-ocean ridges; exponentialdecrease of heat flow orthogonal to spreading centers (Sclaterand Francheteau, 1970); and earthquake distribution at con-vergent margins (Benioff, 1964). Historical accounts of theevolution from a static to dynamic worldview are given byUyeda (1978) and Allegré (1988).

Initial hypotheses of the relationship between differentclasses of ore deposits and their plate tectonic settings wereset out by Rona (1980), Mitchell and Garson (1981), andSawkins (1984). These accounted for the distribution of someore deposit types in the Phanerozoic. However, there werelimitations: (1) at the time, genetic hypotheses for many typesof ore deposit were predicated on syngenesis; (2) where con-sensus existed on a syngenetic versus epigenetic origin, theage of mineralization was not well constrained; (3) epochs, orsecular cycles, of metallogenic provinces were not accounted

for; and (4) extrapolation to the Precambrian met with uncer-tainties as to tectonic processes during that era. Windley(1995) compiled a concise list of metallic and nonmetallic re-sources for each era, documenting their geodynamic and ge-ologic settings.

It is now generally accepted that plate tectonics operatedfrom ~3.4 Ga, albeit in some early form that likely differsfrom today, with intermittently more intense plume activity to1.9 Ga (Fyfe, 1978; Isley and Abbott, 1999). Archean craton-scale faults are commensurate with lithospheric plate interac-tions (Sleep, 1992). In addition, Cenozoic-type convergentmargin arc associations, including the presence of boninites,Mg andesites, and adakites, in Precambrian supracrustal ter-ranes require that arc-trench migration occurred (Polat et al.,2003). An alternative precept of Archean geodynamics isgiven by Hamilton (1998).

Advances in geochronology have resolved many of the un-certainties in the timing of both metal deposits and metallo-genic provinces. This constraint permits evaluation of func-tional relationships between lithotectonic associations,magmatism, pressure-temperature-time (P-T-t) conditionsand fluid compositions, and geodynamic setting, concurrentlyresolving the syngenetic issue (e.g., Kerrich and Cassidy,1994). Based on Meyer’s (1981, 1988) compilations of thespace-time distribution of metallogenic provinces, Barley andGroves (1992) provided insights into the episodic develop-ment of distinct classes of metallic deposits as a function of thesupercontinent cycle. Geologic processes are intrinsically sto-chastic, so there is progressive uncertainty in reconstructingthe supercontinent cycle back through the Precambrian. Yet,this framework confers an elegant account for metallogenicprovinces and their episodicity from 2.7 Ga to the present.

During the last 25 years there have been profound gains inknowledge as to how plate tectonics operates through time,stemming from the heuristic approach of geology as a field andanalytical science. In addition to development of the concept ofthe supercontinent cycle, knowledge has advanced on manyfronts relevant to metal deposits, including: (1) how evolutionof lithospheric mantle controls crustal evolution (Jordan, 1988);(2) recognition of superfamilies of orogenic belts (Sengor andNatal’in, 1996); (3) the role of mantle plumes and their inter-action with lithospheric plates (Condie, 2001; Wyman and Ker-rich, 2002); (4) transitions in both plume and convergent mar-gin magmatism near the Archean-Proterozoic transition(Taylor and McLennan, 1995; Isley and Abbott, 1999); (5) de-velopment of, and processes in, convergent margins (see re-view by Richards, 2003); (6) characterization of geothermal sys-tems on land (Elder, 1981) and submarine counterparts, someof which are actively depositing sulfide minerals, such as in theLau back-arc basin (Ishibashi and Urabe, 1995; Mills and El-derfield, 1995); (7) quantification of global geochemical cycles(Jacobson et al., 2000); (8) seismic tomography (van der Hilstet al., 1998); (9) precise geochronology (Dalrymple, 1991); and(10) the fractal, or scale-invariant, nature of many geologic

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Phanerozoic Pb-Zn SEDEX ores localized in rifted passive continental margins containing evaporites at lowlatitudes. Porphyry Cu and epithermal Au-Ag deposits occur in both intraoceanic and continental margin arcs;ore fluids were related to slab dehydration, peridotite fusion, and hybridization with upper-plate crust. De-posits exposed today are largely <200 m.y.-old, given their low preservation potential in topographically ele-vated ranges.

processes, including of metallogenic provinces (Turcotte, 1992;Weinberg et al., 2004).

Accordingly, in this overview, we reframe the space-timedistribution of ore deposits in terms of four interrelatedprocesses: (1) lithotectonic associations that develop in agiven geodynamic setting, (2) classes of metallogenicprovinces that develop in those associations, (3) secular varia-tions of geodynamic environments in the supercontinent-cycle framework, and (4) secular change of continental lithos-pheric mantle that influences all of the above.

Evolution of near-surface conditions has also been viewedas a control on the distribution of some ore deposits throughtime, specifically those having elements with redox-sensitivesolubility, such as Fe and U. Two polarized schools of thoughtemerged and have persisted. Cloud (1972) proposed a lowpO2 in the Archean, with a transition to oxygenation of Earth’satmosphere-hydrosphere in the Proterozoic, whereas Dim-roth and Kimberly (1976) advocated Archean atmosphericpO2 close to the present atmospheric level (PAL). More re-cently, some workers have promoted the early low pO2 modelbased on mass-independent S isotope fractionation of atmos-pheric S gases, and a rise of atmospheric oxygen at 2.4 to 2.2Ga as the redox state of volcanic gases shifted (Farquar et al.,2000; Holland, 2002). In contrast, Ohmoto maintained thatpO2 was within 50 percent of present atmospheric level by 4Ga, based on Fe mobility in Archean paleosols and on depo-sitional mechanisms for iron formation that are akin to thosepresently occurring in the Red and Black Seas (Ohmoto,1997, 2004a,b). Resolution of this issue is not readilytractable, as many lines of evidence may reflect local condi-tions, and it is difficult to demonstrate preservation of pri-mary signatures (e.g., Clout and Simonson, 2005). It is clearfrom molecular microfossils that the earliest photosynthesisin the Paleoarchean was anoxygenic, using bacteriochloro-phyls, whereas oxygenic photosynthesis by photosystem II, in-volving cyanobacteria, was established by the Mesoarchean(Nisbitt, 2002). This review does not further consider theissue.

No modern text on ore deposits addresses recent advancesin geodynamics. Accordingly, we present a brief synthesis ofgeodynamic concepts as a framework for discussing mineraldeposits. The divisions between geodynamic settings usedhere reflect the preference of the authors. For example, weexplicitly recognize that there is a continuum between domi-nant plume-lithosphere interaction, where magmatic Ni-Cu-PGE deposits form; through belts of anorogenic magmatismthat host Fe-Ti-V deposits, in which plume magmas do notadvect to shallow crustal levels; and to continental rifting withsubdued plume activity, which is the setting for Fe oxide-Cu-Au-REE and sediment-hosted Cu-Co deposits. For eachmain geodynamic setting, we have selected the best charac-terized metallogenic provinces for discussion of the role ofgeodynamics in the formation of a class, or classes, of mineraldeposit, without necessarily including all deposit subtypes.

Geodynamics

Introduction

Plate tectonics is a kinematic theory according to which thelithosphere, the upper layer of the Earth including crust and

lithospheric mantle, is divided into a finite number of plates.The plates are torsionally, but not flexurally, rigid. Plates in-teract at divergent, convergent, and transform-fault bound-aries (Fig. 1A), as they migrate across the surface of the Earth(Isacks et al., 1968; Cox and Hart, 1986; Sengor, 1990). Platemotions are the surface reflection of the fundamental processby which heat is removed from the interior of the Earth.

The oceanic and continental lithospheric plates, alsotermed the mechanical boundary layer (Fig. 2A,B), constitutethe translationally mobile upper boundary layer of the three-dimensional convection cells in the asthenospheric mantle.The core-mantle boundary (referred to as D", 2,900 km deep)is the lower boundary layer of the mantle convection cells.The boundary between upper and lower mantle (D', 670 km)is defined seismologically and reflects a mineralogical phasetransition. The upper and lower mantle probably convects in-dependently, albeit with episodic overturn, based on geo-chemical, heat flow, and seismic evidence (Stein and Hof-mann, 1994; van der Hilst et al., 1998; Butler and Peltier,2002). Heat is removed from the core and mantle to the sur-face by this convection and by plumes that rise from the core-mantle boundary, advecting through the convecting lowerand upper mantle to the surface (Davies, 1999). Heat passesfrom the convecting asthenospheric mantle through the tor-sionally rigid lithospheric plates either by conduction or byadvection of magmas. Thermal boundary layers form at thetransition from convecting to convecting or convecting toconducting domains; they are present at the D" core-mantleboundary, at the D' upper-lower mantle transition, and be-tween the base of the lithosphere and top of the convectingupper mantle, which is also the low-velocity zone (Fig. 2).

Subducting oceanic lithospheric plates penetrate the D'upper-lower mantle boundary at 670 km, as imaged by seis-mic tomography, and probably are stored in lithosphericgraveyards at the core-mantle boundary (D"), where they aresporadically reactivated as mantle plumes. Similarly, ananomalously hot mantle plume, extending into the lowermantle, has been imaged beneath the Iceland ocean plateau(Bijward and Spakman, 1999; Kárason and van der Hilst,2000). Accordingly, there is mass as well as heat exchange be-tween the upper and lower mantles.

Oceanic and continental lithosphere

Schematic diagrams depicting the tectonic setting of oredeposits generally stop at the base of the deposit or the crust,the petrological seismic Mohorovic discontinuity (Moho).However, the larger context in which mineral concentrations,i.e., deposits, form should more comprehensively be consid-ered in a lithosphere-asthenosphere framework that reflectsgeodynamic settings. These in turn control the conjunction ofstructures, magma reservoirs, fluid reservoirs, basins, andtheir interactions (Fig. 2).

Modern oceanic lithosphere has a ~6-km-thick basalticcrust and ~30- to 50-km-thick lherzolitic mantle lithospherenear ridges. The mantle lithosphere thickens to a maximumof ~70 to 100 km as it progressively cools with increasing dis-tance from the oceanic spreading axis, by accretion of under-lying asthenosphere (Fig. 2A; Keary and Vine, 1996). Com-pared to the underlying asthenosphere, oceanic lithosphereaway from ridges is relatively cool, mechanically rigid, and

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FIG. 1. A. Map of continental and oceanic lithospheric plates. Triangles signify polarity of subduction, trenches migratein the opposite direction as slabs sink approximately vertically. Length of arrows proportional to plate velocity. Red symbols= Cordilleran superfamily of orogenic belts; green symbols = continent-continent superfamily of orogenic belts. Modifiedfrom Condie (1997). B. Distribution of Archean cratons and Proterozoic and Phanerozoic terranes. After Kusky and Polat(1999). C. Thickness of continental lithospheric mantle from Artemieva and Mooney (2001).

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FIG. 2. A. Cross section through oceanic lithosphere, modified from Keary and Vine (1996). B. Cross section through con-tinental lithosphere, illustrating the thick, refractory irregular base or keel of the continental lithospheric mantle, distinctiveof Archean cratons. This mantle includes subcreted plateau lithosphere metasomatized by subduction at shallower levels, thesource of Neoarchean and Proterozoic cratonic norites. Deeper levels are the residue of plume melting, buoyantly coupledto overlying continental lithospheric mantle and crust. Such Archean mantle is refractory and thus is responsible for the highpreservation potential of Archean mineral deposits; this level includes the diamond facies. Translithospheric structures arefocused at the transition to thinner Proterozoic and younger continental lithospheric mantle, controlling the location ofplume-related Ni-Cu-PGE and Fe oxide-Cu-Au-REE deposits. Modified from Nixon and Davies (1987), Artemieva andMooney (2001), and Wyman and Kerrich (2002). C. Cross section through oceanic crust, illustrating the location of VMS de-posits that form in back arcs and podiform Cr deposits generated at intraoceanic suprasubduction zones. Modified fromKeary and Vine (1996). D. Age-thickness relationship of continental lithospheric mantle from velocity structure (afterArtemieva and Mooney, 2001). E. Depth-differential strength relationships of oceanic and continental lithosphere; foroceanic lithosphere this relationship controls the thickness of obducted ophiolites; for continental lithosphere the minimumat ~35 km controls the thickness of accreted terranes. F. Depth-shear wave velocity relationships of different geodynamicsettings. (E) and (F) modified from Keary and Vine (1996).

negatively buoyant (Fig. 2A). For a hotter Archean uppermantle, greater degrees of melting occurred at spreading cen-ters (Bickle, 1986). According to calculations of Abbott et al.(1994a), Neoarchean basaltic oceanic crust was ~11 km thick,with a commensurately thicker mantle lithosphere residuedepleted in incompatible elements from basalt extraction.Consequently, Archean ocean lithosphere would have sub-ducted at shallower angles from thermal and buoyancy con-siderations. There has been a secular decrease in the temper-ature of mantle plumes; accordingly, ocean plateau crust hasalso become thinner through time (Fig. 3).

The continental lithosphere has a 30- to 80-km-thick crustalsector in Archean and younger eons. Continental lithosphericmantle is 250 to 350 km thick under Archean continentalcrust but ~150 km thick for Proterozoic and ~100 km for

younger terranes. The existence of Archean continentallithospheric mantle defines cratons (Figs. 1B,C, 2 B, D, F;Artemieva and Mooney, 2001; Plomerova et al., 2002), and itis more buoyant and refractory than mantle lithosphere be-neath younger continental regions; its thickness and thermalstructure lead to the preservation of diamonds (Fig. 2B,C).Archean continental lithospheric mantle is the residue ofdeep melting in hot plumes coupled to crust. There is a bi-modal depth distribution to such mantle at 350 to 300 km andat 220 to 200 km, with the former characterizing blocks >6 to8 × 106 km2 in area (Artemieva and Mooney, 2001), and withimplications for diamond potential. Younger plumes were lessfrequent and cooler, so they did not generate refractoryresidues (Fig. 2; White, 1988; Jordan, 1988; Pollack, 1997;Herzberg, 1999; Artemieva and Mooney, 2001). For example,the continental lithospheric mantle is 190 to 240 km thick inthe diamondiferous Magan and Anabar cratons but thins to150 to 180 km for the Proterozoic Olenek province (Griffin etal., 1999). From studies of xenolith suites, there is a seculartrend from highly depleted harzburgites in Archean conti-nental lithospheric mantle, through intermediate depletion inthe Proterozoic, to mildly depleted lherzolites in the Phaner-zoic. Archean continental lithospheric mantle has a density of3.36 g/cm3, whereas Proterozoic continental lithosphericmantle is 3.38 g/cm3, marginally less dense than ambient as-thenosphere (Griffin et al., 2003).

Archean supracrustal terranes are dominated by bimodalvolcanic arc sequences and postvolcanic tonalite-trond-hjemite-granodiorite batholiths, whereas Archean continentallithospheric mantle is refractory harzburgite, with the com-position of the residue of plume melting. This apparent para-dox may be resolved if migrating arcs captured ocean plateauserupted from mantle plumes. Buoyant plateaus jam subduc-tion zones, generating composite arc-plume crust, and thebuoyant residue of plume melting couples to the base of thecrust (Wyman and Kerrich, 2002). Prior to capture and cou-pling of plume residue, subduction caused metasomatism ofperidotitic subarc lithosphere. During subsequent exten-sional events, and/or plume impingement, metasomatized do-mains melted to generate the voluminous noritic magmascharacteristic of Neoarchean to Proterozoic layered igneouscomplexes in or near Archean cratons (Fig. 2B; Hall andHughes, 1980). Those magmas are integral to formation ofNi-Cu-PGE and Fe-Ti-V deposits. Proterozoic and youngerplumes were not hot enough to generate refractory residue;consequently, Proterozoic and younger continental lithos-pheric mantle is thinner, denser, and less refractory, such thatcrustal terranes are more readily reworked during subsequentorogenies (Figs. 1, 2).

During collisional orogens in the Proterozoic and Phanero-zoic, both crust and continental lithospheric mantle thicken,and part of the latter may delaminate; hot asthenosphere thenflows under thinned lithosphere, creating elevated orogens,as in the Tibetan plateau (Houseman and Molnar, 1997).During lithosphere thickening under compression, radioac-tive heat weakens the crust, and decoupling of lower crustand continental lithospheric mantle may occur at the base ofthe upper felsic crust (Meissner and Mooney, 1998). High-temperature–low-pressure metamorphism and extensionalcollapse with escape tectonics ensue, in conjunction with

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FIG. 3. A. Plume intensity through time, simplified from Abbott et al.(1994a). Ocean crust production, dashed line. B. Sea level change throughtime; the first extensive exposure of continents above sea level occurred afterthe 2.8 to 2.6 Ga plume maxima, followed by development of extensive pas-sive margin sequences at ~2.4 to 2.2 Ga as the supercontinent Kenorland dis-persed. C. Decrease of potential shallow mantle temperature. D. Decreasein thickness of ocean crust in response to secular change of shallow mantletemperature (Abbott et al., 1994) and of plateau crust in response to chang-ing plume temperature.

asthenospheric and crustal magmatism. Delaminated conti-nental lithospheric mantle has been imaged by teleseismic to-mography beneath the Alpine-Himalayan orogen (Schott andSchmeling, 1998). Delamination is in progress beneath theBasin and Range province and Tibetan plateau, is interpretedto have occurred beneath the Puna plateau of northwesternArgentina (Kay and Kay, 1993), and characterized the latestages in the development of the Variscan and Grenvilliancontinent-continent orogens (Windley, 1995).

The low-velocity zone is the thermal boundary layer be-tween torsionally rigid lithospheric plates and the convectingasthenosphere; low S wave velocities result from domains ofpartially melted lherzolite, conferring low strength. This zoneis 100 to 200 km thick below ridges where thermal gradientsare high, thinner below normal continental lithosphere, and isthin to absent beneath Archean continental lithospheric man-tle where thermal gradients are low (Fig. 2 A, B; Keary andVine, 1996).

Characteristics of plate boundaries

Divergent plate boundaries: As oceanic plates separate atridges due to far-field extensional forces, decompressional melt-ing of asthenospheric mantle generates mafic magmas that ac-crete to the edges of plates to form new crust (Keary and Vine,1996). Upwelling of asthenospheric upper mantle beneathridges is passive, in response to plate separation. In a simplifiedcross section, the oceanic lithosphere is composed of lower ul-tramafic mantle (mantle tectonites, dunites, lherzolites, andharzburgites) at the base, and mafic crustal rocks (gabbros,sheeted dike complex, and basalts) at the top, bounded by theoceanic Moho. The thickness of the lithosphere increases fromzero at ridges to 70 to 100 km at an age of ~70 m.y., then main-tains approximately uniform thickness, as plates move awayfrom spreading centers. Commensurately, the depth of theocean floor increases with the age of oceanic lithosphere, due tothermal cooling of the lithosphere associated with thickeningand subsidence (Fig. 2A; Parsons and Sclater, 1977).

Convergent plate boundaries: At convergent margins, theplate with higher density sinks beneath the lighter plate,forming a subduction zone, and the leading edge of the over-riding plate becomes a paired fore arc and magmatic arc.Where two oceanic plates converge, the older and denseroceanic plate generally sinks beneath the younger and lighterone, generating oceanic island arcs, such as the Marianas andthe south Sandwich arcs. Given its higher density, oceaniclithosphere subducts underneath continental lithosphere toform a continental magmatic arc, such as the Andean, Suma-tran, and Japanese arcs.

Convergent margins generally feature the following tec-tonic elements: (1) a deep marine trench seaward of the forearc; (2) a subduction-accretion complex located between theunderriding plate and the fore-arc basin; (3) a fore-arc basinbetween the arc axis and the subduction-accretion complex;(4) a magmatic arc; and (5) an inboard foreland basin-thrustbelt, which undergoes subsidence and sedimentation due totectonic loading, tectonic imbrication, and later compression-driven uplift (Fig. 4). Porphyry Cu deposits form in oceanicand continental arcs, and most preserved volcanic rock-asso-ciated massive sulfide deposits form in oceanic arcs oroceanic or continental back arcs.

Based on relative plate motions, magmatic arcs are dividedinto extensional, neutral, and compressional (Dewey, 1980;Sengor, 1990). Extensional arcs, such as the Marianas, arecharacterized by dominantly mafic volcanism, back-arc basinopening, an ophiolitic fore-arc basement, deep trenches, andsteeply dipping Wadati-Benioff zones. Given its thermallyweak nature, arc lithosphere generally undergoes extension toform an intra-arc basin or an intra-arc spreading center; theLesser Antilles and Taupo arcs are examples of initial stages,whereas the Lau basin has evolved into a back arc.

Compressional arcs, such as the Central Andes, lie on con-tinental lithosphere, and are characterized by mainly inter-mediate to felsic magmatism, back-arc thrusting, continentalfore-arc basement, shallow trenches, and shallow Benioffzones. Neutral arcs such as the Central American, Sumatran,and Alaska Range-Aleutian arcs have characteristics interme-diate between extensional and compressional arcs and usuallyhave large subduction-accretion complexes and orogen-paral-lel strike-slip faults (Windley, 1995).

Arc magmatism varies along and across strike. All arc mag-mas are characterized by variably light rare earth element(REE) and lithophile element (Cs, Rb, Ba, K, and Pb) earthelement enriched patterns and depletions in Nb, Ta, P, and Ti(Pearce, 1982; Saunders et al., 1991; Keleman et al., 2004).Tholeiitic magmatism is dominant between the fore-arc basinand arc axis; calc-alkaline magmatism occurs mainly in thecentral region of the arc, whereas late alkaline igneous rockstend to occur between the arc axis and back-arc region (theK-h relationship; see Wilson, 1989, for a review). The com-position of continental crust requires that mafic cumulatesfounder under arc crust (Rudnick and Gao, 2004), with spaceconservation accommodated by inflowing asthenosphere. Re-gional metamorphism varies from subgreenschist to eclogitefacies (Fyfe et al., 1978), and the occurrence of adjacent high-temperature and/or low-pressure (greenschist) and high-pressure and/or low-temperature (blueschist) metamorphicbelts is unique to convergent plate boundaries (Ernst, 1975).

The uppermost section of subducting oceanic lithosphere isprevalently marine turbidites but may include pelagic sedi-ments, oceanic islands, seamounts, and carbonate platforms.These are commonly scraped off, deformed, metamorphosed,and accreted to the base of the overriding plate to form a sub-duction-accretion complex. Complex interaction betweenoverriding and subducting plates results in thrusting, folding,and mélange formation within the subduction-accretion com-plex, with late transpression and associated strike-slip fault-ing. In arcs characterized by strong coupling between theoverriding and subducting plates, attrition of the fore arc oc-curs by subduction-erosion (von Huene et al., 2004). Trenchturbidites have a catchment in the upper levels of subduction-accretion complexes. Plate movement is driven by the nega-tive buoyancy of subducting slabs, not by mantle convection(Conrad and Lithgow-Bertelloni, 2002). Stern (2002) has re-cently reviewed processes in subduction zones.

Transform plate boundaries: Transform, or conservative,boundaries accommodate the motion from divergent- to con-vergent-plate boundaries and accommodate translation be-tween ridge sectors spreading at different rates, as requiredby plate motion on a spherical surface (Wilson, 1965). Trans-form-plate boundaries separating continental lithospheric

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FIG. 4. A. Life span-geodynamic relationships of sedimentary basins. Modified from Woodcock (2004). Abbreviations: BA= back arc, FA = fore arc, FL = foreland, IA = intra-arc, O = oceanic, PM = passive margin, R = continental margin rift, RA= retro-arc, SS = strike slip, T = trench, TS = trench slope. (A) after Kyser et al. (2000), (B), (C), and (D) modified from Ross(2000), (E) a composite from miscellaneous sources and R. Kerrich.

blocks are termed transcurrent or continental strike-slipfaults. Examples are the San Andreas fault zone of California,the North Anatolian strike-slip fault zone in Turkey, and theTintina and Denali fault zones of western Canada and Alaska.Transtensional regions are characterized by normal faulting,pull-apart basins, and dominantly basaltic volcanism, whereastranspressional regions feature thrusting, folding, and uplift,in addition to strike-slip faulting in both cases (Christie-Blickand Biddle, 1985; Sylvester, 1988).

Superfamilies of orogens

Cordilleran orogens: Sengor and Natal’in (1996a) classifiedorogenic belts into two superfamilies, Cordilleran and conti-nent-continent. This insight has profound implications formetallogeny. Cordilleran-type orogens, also referred to asTurkic or transpressional (Sengor and Natal’in, 1996a), exter-nal (Murphy and Nance, 1992), or accretionary (Windley,1995), represent continental growth via the process of terraneaccretion. These sutured tectonostratigraphic terranes arefragments of juvenile arcs and ocean plateaus, plus marinesedimentary rocks, tectonically assembled in accretionaryprisms; there is typically little addition or reworking of oldercontinental crust (Ben-Avraham et al., 1981). Collision of ter-ranes occurs dominantly in an oblique manner, with the par-titioned compressional component responsible for much ofthe orogeny. Cordilleran-type orogens are characterized byboth extensive lateral and vertical accretion above a subduct-ing slab. Where subduction-erosion hinders terrane collisionand thus lateral accretion, Andean-type orogens dominate.These possess the arc-related porphyry and epithermal de-posits that also characterize Cordilleran-type orogens but lackmost other, more deeply formed deposit types that are com-mon throughout the blocks of allochthonous juvenile crustwithin such orogens, as described below.

Multiple sutures at terrane boundaries are inherent tolong-lived terrane collison. Sutures commonly serve as sitesfor ensuing economic mineralization. Seaward growth of con-tinental margins, with such sutures defining progressive ter-rane accretion, tends to be a long-lived process of perhaps~300 to 400 m.y.; examples include the Cordilleran orogen,370 Ma to present; Altaid orogen, 610 to 250 Ma; and Pan-African orogen, 900 to 630 Ma (Burchfiel et al., 1992; Sengorand Natal’in, 1996a). Depending on the degree of obliquity toeach terrane collision, these sutures behave as thrust and/orstrike-slip faults. In many cases, lateral displacement of ter-ranes becomes relatively more common late during orogene-sis or even subsequent to all collision. Such transform conti-nental margins concentrate juvenile crust, and likelyassociated mineral deposits, in restricted regions of an evolv-ing orogen (Patchett and Chase, 2002). The terranes of theAltaid orogen underwent thousands of kilometers of left lat-eral and right lateral movements during the final stages of Pa-leozoic tectonism in central Asia (Sengor and Natal’in,1996a). Major shifts from compressional to more translationalregional stress regimes appear conducive to seismic eventsand extensive episodes of fluid flow (Sibson et al., 1988; Ker-rich and Wyman, 1990) and may be important controls on thedevelopment of large orogenic gold provinces in Cordilleranorogenic belts (e.g., Goldfarb et al., 1991, 2005). Given far-field compressional regimes superimposed on more localized

transtensional to transpressional zones late during orogenesis,Cordilleran orogens generally undergo significant oroclinalbending (e.g., Alaska and the Altaids; Yakubchuk et al., 2002,2005). These strike-slip regimes also cause the highly dis-membered nature of ophiolite sequences within most oro-gens and thus a discontinuous distribution to many preaccre-tionary VMS and chromite ores.

Cordilleran-style orogens may show a similarly wide(>1,000 km) pattern of subduction-related magmatism, as inthe Altaids and mainland Alaska. By contrast, continent-con-tinent orogens feature narrow magmatic arcs. In Cordilleanorogens, the ages of the igneous rocks young toward theocean, as arc magmatism migrates episodically in that direc-tion as the continental margin is built outward (Sengor andNatal’in, 1996a). Ages of orogenic gold deposits tend to followthe same approximate spatial and/or temporal pattern (Gold-farb et al., 1997). Most igneous rocks are juvenile in the oro-gens; there are limited examples of remelted crust seaward ofthe craton edges (Windley, 1995).

Lithological units in Phanerozoic orogens are dominated bydeep marine turbidite sequences and lesser basalts andcherts, as these are the dominant rocks being accreted off thetop of subducting oceanic slabs and comprising the growingprism defining the arc-trench gap (Fig. 4E). Addition of newcrust to a craton margin is also common in a spreading back-arc regime, where foreland or retroarc basins may evolve in aregion of extension between the continental arc and cratonedge. These units may have a higher volume of fine-grainedterrigenous and biogenic material than units in the fore arc,which are more likely dominated by the clastic products ofdeep-sea turbiditic currents. Importantly, the pelitic sedi-mentary rocks and related mafic volcanic and volcaniclasticsequences commonly contain volatile-rich mineral phasessuch as phengite, biotite, lawsonite, chlorite, dolomite, mag-nesite, and pyrite, all potential contributors of H2O, CO2, andS to fluid phases produced during later thermal and de-volatilization events (e.g., Fyfe et al., 1978). Furthermore,marine pyrite may contain trace amounts of gold that can alsobe mobilized during subsequent heating of the marine rocks.

During the last decade, Cordilleran-style orogens, as prod-ucts of a present-day style of plate tectonics, have becomewidely accepted as having developed far back into the Pre-cambrian (Sengor and Natal’in, 1996a; deWit, 1998). Archeanand Paleoproterozoic terranes are dominated by greenstonesand tonalites, with minor turbidites; these linear belts alsolikely formed via processes of accretion at convergent plateboundaries (Kusky and Kidd, 1992; Kusky and Polat, 1999;Foley et al., 2002). Structural style and metamorphic regimesin many Precambrian greenstone belts, or compositetectonostratigraphic terranes, resemble those in PhanerozoicCordilleran-style orogens. For example, accretionary assem-bly of the ~2.7 Ga Superior province and Yilgarn cratons(Sengor and Natal’in, 2004), including the involvement of sig-nificant volumes of plume-derived oceanic plateau crust(Stein and Hofmann, 1994; Polat et al., 1999), was character-ized by terrane accretion and batholith emplacement that mi-grated in a seaward direction (Kerrich et al., 2000).

Ophiolites, long appreciated as footprints of Cordilleran-style tectonics, are now widely recognized in Precambrian en-vironments, with a particularly high abundance of tholeiitic

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pillow basalts in many cratons (de Wit, 2004). Higher geot-herms in the Archean are reflected by the widespread high-grade gneissic basement rocks, which, with refractory conti-nental lithospheric mantle, have preserved the mid-crustalCordilleran-like greenstone belts for billions of years. Simi-larity in the geologic evolution of Precambrian and Phanero-zoic Cordilleran-style continental margins is reflected in asimilar metallogenic record being preserved in metamor-phosed rocks of all such orogens, regardless of geologic age(Goldfarb et al., 2001).

Continent-continent orogens: A second type of orogen istermed continent-continent collisional or Tethyan. It is typi-cally marked by the closure of an ocean basin, a single well-defined Z- or C-shaped suture zone containing ophiolites be-tween blocks of continental crust, a magmatic arc on theactive margin, and deformation of passive margin sequences.Collision is orthogonal to oblique, with an exceptional amountof crustal thickening, and reworking of the older crustalblocks (Windley, 1995; Sengor and Natal’in, 1996a). This tec-tonism includes metamorphism, widespread partial meltingof the lower crust during lithosphere thickening, delamina-tion, and commonly underplating by mafic magmas. Depend-ing on the structural complexity, these orogens may showabundant, high-level overthrusting exemplified by the Alpinetype or limited thrusting of allochthonous blocks as in the Hi-malayan type (Sengor, 1990; Sengor and Natal’in, 1996a).

Mantle plumes

Pirajno (2000) gives a comprehensive treatment of mantleplumes and ore deposits upon which this section draws ex-tensively. Jets of anomalously hot mantle are ejected fromthermal boundary layers, most likely the core-mantle bound-ary at 2,900 km, which advect through the mantle by thermalbuoyancy on timescales of only 10 to 50 m.y. The plume headis 500 to 1,000 km in diameter, whereas the tail, which feedsthe head, is ~100 km in diameter. At the top of the uppermantle, ambient temperature is ~1,280ºC, the plume head~1,480ºC, and the tail ~1,700ºC. Plumes conductively heatambient mantle, which is entrained into the plume head. Onimpinging upon normal lithosphere at ~150-km depth, theplume head flattens to 1,000 to 2,000 km while undergoingextensive decompressional melting (White, 1992). Anom-alously hot plumes, with high buoyancy-driven flux, advectbasalts through continental lithosphere to erupt as continen-tal flood basalts. Basaltic liquids from cooler plumes, or fromadiabatically decompressed asthenosphere under thinnedcontinental crust, pond at the Moho density filter (Herzberget al., 1983); here they fractionate to form anorogenic gabbro-anorthosite complexes that may host Fe-Ti-V deposits(Cawthorn et al., 2005) and also fuse refractory lower crustinto A-type granites with which Fe oxide-Cu-Au-REEprovinces are associated (Fig. 2B; Windley, 1995; Williams etal., 2005).

Crucial to the understanding of magmatic Ni-Cu (Arndt etal., 2005; Barnes and Lightfoot, 2005) and chromite deposits(Cawthorn et al., 2005), as well as deposits associated withanorogenic magmatism, is that plumes do not melt by de-compression at ~250 km beneath Archean continental lithos-pheric mantle but rather penetrate laterally as dikes. Theseinclude the 2596 Ma Great Dyke and 2200 Ma Matachewan

swarm. Alternatively, the plumes spread laterally under thenormal continental lithospheric mantle (Fig. 2B). Plume ac-tivity, particularly of superplumes, is episodic, with maxima at~3.8, 3.4, 3.0, 2.7, 2.4, 1.9, and 1.7 Ga, with one at ~250 Maand another superplume in the Cretaceous (Fig. 3A; Larson,1991; Ernst and Buchan, 2001; Abbott and Isley, 2002).

Mantle plumes occur in three broad varieties, as discussedbelow.

Long-lived hotspots with low magma flux: These plumesgenerate ocean islands, such as the Emperor-Hawaii chain onoceanic lithosphere, or hotspot tracks on continents, e.g., theColumbia River-Yellowstone track spanning 45 Ma to the pre-sent (Schissel and Smail, 2001).

Short-lived plumes that generate flood basalt provinces:Plumes that erupt through oceanic lithosphere form oceanicplateaus, including Kerguelen, Ontong-Java, and Iceland, orcontinental flood basalts. For the Siberian and Deccan conti-nental flood basalts, 1 to 3 × 106 km3 of flows erupted during<1 m.y; tholeiitic basalts predominate, with minor alkalibasalts and picrites. The Tertiary North Atlantic igneousprovince, which includes continental flood basalts on Green-land, a volcanic passive margin on eastern North America,and the Iceland plume, collectively represent a transitionfrom continental flood basalts to an ocean plateau as NorthAmerica and Scandinavia rifted apart. The three elements ofsuperplumes, continental flood basalts, giant dike swarms,and mafic intrusive complexes, are collectively referred to aslarge igneous provinces (Coffin and Eldholm, 1994; Saunderset al., 1997; Eldholm and Coffin, 2000). All three elementsare present in the 1267 Ma Mackenzie giant dike swarm,Coppermine CFB, and the Muskox intrusive complex ofnorthern Canada (Ernst and Buchan, 2004). Giant dikeswarms may represent a failed triple junction and, therefore,point toward a paleo-ocean (Fahrig, 1987). The secular distri-bution of iron formations, from 3.8 Ga to 40 Ma, is controlledby mantle plumes.

Superswells or mantle upwellings: These features have di-ameters of ~10,000 km and spawn hotspots. There are twoknown, one centered on the South Pacific and another belowAfrica, both with dynamic topography (McNutt, 1998). TheAfrican superswell was responsible for rifting of Gondwanafrom Laurasia.

Plumes and ore deposits: All three expressions of mantleplumes have a role in mineral provinces, from diamond fields(Gurney et al., 2005) to Ni-Cu-PGE deposits (Pirajno, 2000).

Hotspot plumes are approximately fixed with respect to themantle, so ocean island chains provide a reference frame forhotspots that constrains plate motions (Norton, 2000). Mantleplumes and lithospheric plate motions are not stronglycoupled. However, where a plume erupts proximal to a spread-ing center it may capture the ridge, as with the Icelandplume–Mid-Atlantic Ridge. Plumes may interact with conver-gent margins, such as impingement of the mid-CretaceousMarie Byrd Land plume with the Phoenix plate subducting be-neath Antarctica (Weaver et al., 1994). Present-day examplesinclude the Samoan plume proximal to the Tonga trench andinteraction of the Yellowstone plume with the Farallon plate(Schissel and Smail, 2001). Plume-ocean ridge and plume-con-vergent margin interactions cause some of the largest knownstructural and geochemical anomalies (Ito et al., 2003).

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Ocean plateaus with >30-km thickness of basaltic crust,erupted from anomalously hot mantle plumes, resist subduc-tion, and cause collisional orogenesis when they jam upagainst a subduction zone (Cloos, 1993). The Solomon-NewIreland arc has migrated to capture the 120 to 90 Ma Ontong-Java ocean plateau, which is being jammed against the sub-duction zone; this is where the Lihir Au deposit has formed(MacInnes et al., 1999). Formation of the giant 2.7 Ga KiddCreek VMS deposit followed capture of the Abitibi arc by anocean plateau (Wyman et al., 1999)

There is compelling evidence for the influence of mantleplumes on conditions of surface geology, the hydrosphere, at-mosphere, and biosphere (Larson, 1991; Coffin and Eldholm,1994; Kerr, 1998). Isley and Abbott (1999) and Condie et al.(2001) demonstrated a coincidence in timing of mantleplumes, deposition of iron formations and black shales, andthe chemical index of alteration. Ocean plateaus that eruptedfrom plumes formed thick crust that displaced oceans acrosscontinents and caused flooding of continental shelves; theplumes also resulted in the discharge of Fe-rich hydrothermalfluids and the release of CO2 and other gases that generatedgreenhouse conditions, causing intense silicate weathering(Kerr, 1998).

Sedimentary basins

The geodynamic setting of sedimentary basins, and theirlifespan and fate, have been summarized by Ross (2000) andWoodcock (2004). This discussion deals only with foreland,intracontinental, passive margin, and oceanic basins, drawingmainly on these summaries (Fig. 4).

Foreland basins develop as a consequence of tectonic loadingat convergent margins. A classic profile involves a foredeep axisproximal to an orogen, a continental ramp or outer slope, and aperipheral bulge. Lithosphere elastic thickness determinesbasin characteristics; transitions from narrow, deep-water flyschsequences to wide, marine or fluvial molasses facies reflectpropagation of the load from elastically thin lithosphere at a sea-ward position to thicker continental lithosphere. Proterozoicunconformity U deposits and Phanerozoic Mississippi Valley-type (MVT) Pb-Zn deposits accumulated in foreland basins thatevolved to intracratonic basins (Fig. 4B, D).

The pattern of stratigraphic onlap (so-called steersheadgeometry) of intracratonic and passive margin sequences isconsistent with extension being driven by far-field forces, inwhich differential tensile strength causes mantle lithosphereto extend over a wider area than the crust (Fig. 4B,C; Whiteand McKenzie, 1988). The Williston as well as Michigan andIllinois basins developed inboard of the Cordilleran and Ap-palachian orogens, respectively, but the cause of this relation-ship is not clear (Ross, 2000). According to Pysklywec andMitrovica (2000), some intracratonic basins stem from dy-namic topography generated by foundering of subductedlithosphere. Sublithospheric loading generates flexural wave-lengths one order of magnitude longer than surface loads, ac-counting for both the relative dimensions and lifespans of in-tracontinental versus foreland basins (cf. Woodcock, 2004).Proterozoic sedimentary-hosted SEDEX Pb-Zn deposits de-veloped in intracontinental rifts (Leach et al., 2005a,b).

Passive-margin sequences that develop as intracontinentalrifts evolve into ocean basins. A typical sequence is rifting of

continental lithosphere followed by sedimentation, magma-tism linked to thinned continental lithosphere, and evolu-tion to ocean lithosphere. The Atlantic margin, with its con-tinental shelf, continental slope, and rise, is a typicalexample. The sedimentary wedge may be deposited at nor-mal, oblique, or transform continental margins. Transferfaults accommodate differential extension rates and patternsof sedimentation. Subsidence initiates by lithospheric thin-ning from far-field forces and then evolves by thermal con-traction and sediment loading. Basins driven mainly by ther-mal subsidence are characterized by concave-up subsidencepatterns, as documented for aging oceanic lithosphere,whereas foreland basins have concave-down subsidence pat-terns (Fig. 4C; Ross, 2000).

Phosphorites and iron formations accumulated on passivemargins from ~2.4 Ga. Rifted passive-margin clastic sedi-mentary sequences, formed at low latitudes, are favorablehosts for Phanerozoic Pb-Zn ores. The deposits are generatedby metal-rich brines that evolved in adjacent carbonate unitsand basement (Leach et al., 2005a,b). Placer deposits of Ti-Zr-Hf are preserved in Teriary and younger passive marginsequences (Freeman and Donaldson, 2004).

Where extension is focused within a continent, as in theBasin and Range province, a continental back-arc basin maydevelop. The Bathurst and Iberian pyrite VMS provinces areexamples of continental back-arc basins that closed; sill-sedi-ment complexes in the Gulf of Cortez may be a present-dayanalog (Boulter, 1993).

The supercontinent and/or superevent cycle

The concept of the supercontinent cycle emerged in thelate 1980s from recognition that the continental masses as-semble and disaggregate in a cyclic pattern on a timescale of200 to 500 m.y. (Fig. 5; Hoffman, 1988; Murphy and Nance,1992; Rogers, 1996; Rogers and Santosh, 2004). All of thepresent continents formed a single landmass, Pangea, thatbroke up ~180 Ma. Previous supercontinents were Kenorlandat ~2.7 to 2.2 Ga, Columbia at ~1.7 to 1.4 Ga, and Rodinia at~1.0 at 0.6 Ga (Fig. 5; Condie, 2004; Zhao et al., 2004).

A consensus has emerged that rifting of continents and dis-persal of supercontinents is generally triggered by a mantleplume, in keeping with Ziegler’s (1993) estimates of tractionalforces for plumes that impinge on continents (White, 1992;Duncan and Turcotte, 1994; Carlson, 1997). Sill-sedimentcomplexes of the Mesoproterozoic Sullivan Pb-Zn deposit andNeoproterozoic basalt sequences associated with the CentralAfrican Cu province are expressions of mantle plumes that dis-persed the supercontinents Columbia and Rodinia, respec-tively. Condie (1998, 2004) envisaged superevent cycles at 2.7,1.9, and 1.2 Ga in which graveyards of subducted oceaniclithosphere, stored at the 670-km D' boundary, avalanched tothe core-mantle boundary, thus ejecting plumes from thatboundary and causing plume bombardment under the lithos-phere (Fig. 5). Larson (1991) associated the increased rate ofocean crust formation at ridges and plateaus in the PacificOcean with a superplume ejected from the core-mantleboundary, coinciding with cessation of magnetic field reversalsat 41 Ma (for a contrary view see Anderson, 1994).

Murphy and Nance (1992) recognized two principal stylesof supercontinent aggregation, which they termed internal

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FIG. 5. A. Secular distribution of collisional orogens and juvenile crust, with supercontinents (modified from Condie,1997; Columbia after Zhao et al., 2004). B. Secular distribution of mineral deposits, modified from Meyer (1988). C. Super-continent cycle, modified from Rodgers (1996).

and external. Internal aggregation corresponds to continent-continent collision, for exmple, the Alpine-Himalayan, Ap-palachian, and Grenville orogenic belts. External aggregationcorresponds to Cordilleran-style tectonics, where allochtho-nous tectonostratigraphic terranes are transpressively ac-creted to a continental margin. Neoarchean magmatic-accre-tionary events in the Superior and Slave provinces of Canada,Finland, southern Africa, India, and Western Australia likelycorrespond to an early external supercontinent aggregationthat was associated with development of orogenic goldprovinces (Kerrich and Wyman, 1994). Internal cycles involveinternal oceans between continents. The North and South At-lantic Oceans have opened and closed two or three times, asNorth America-South America and Europe-Africa divergedand then closed in Wilson cycles. The Pacific Ocean is an ex-ternal ocean outboard of the external Cordilleran orogen.

Supercontinents may assemble in two configurations. In-troversion involves breakup, opening then closing of interioroceans, and reassembly. In extroversion, following supercon-tinent dispersal, exterior margins of continental fragments ro-tate and collide during reassembly. Combinations of theprocesses may occur. The Paleozoic Appalachian-Caledonian-Variscan orogen is an example of supercontinent introversion.In contrast, during the Neoproterozoic East African andBrasiliano orogens, the exterior ocean surrounding Rodinia,which broke up at ~750 Ma, was consumed during the amal-gamation of Gondwana, representing extroversion (Murphyand Nance, 2003).

Metallogenic provinces in a supercontinent cycle framework

In an important synthesis for economic geology, Barley andGroves (1992) showed that the temporal distribution of sev-eral major classes of metallic mineral deposits can be relatedto the cyclic aggregation and breakup of the continents in thesupercontinent cycle. Metal deposits related to continentalrifting (sedimentary rock-hosted Cu and Pb) would formmainly during initiation of supercontinent fragmentation,whereas deposits related to convergent tectonics (porphyryCu, VMS, orogenic Au) predominate during periods of sub-duction and supercontinent aggregation (Fig. 5).

Superimposed on this ~500-m.y. cycle are variations aris-ing from preservation, thermal decay, and subtleties of tec-tonic style. The scarcity of porphyry Cu and epithermal Audeposits in rocks older than 200 Ma is widely considered tobe the consequence of their low preservation potential inrapidly eroded magmatic arcs and collisional mountainbelts. Preservation potential is considered to be higher inexternal (Cordilleran style) than internal (continent-conti-nent) mountain belts (Barley and Groves, 1992). Thechange in style of base metal-bearing VMS deposits, fromArchean Abitibi type to the Phanerozoic Kuroko and Cyprustypes, may reflect differences in style of subduction, natureof the mantle wedge, and composition of arc magmas, andthese differences in turn stem from decreasing thermal gra-dients. Archean crust is resistant to reworking in youngerorogenic events due to its thick, refractory continentallithospheric mantle. This characteristic accounts for preser-vation of the prodigiously rich orogenic gold provinces ofNeoarchean greenstone terranes (Cordilleran-type accre-tion), VMS (back-arc) camps of the Superior province, and

komatiite-associated Ni deposits (Figs. 1, 2, 3, 5; Kerrich etal., 2000; Groves et al., 2005).

The abundance of VMS deposits in the Superior province,particularly when compared to the sparseness of similar de-posits in Neoarchean counterpart terranes of India, southernAfrica, and Western Australia, might be considered contra-dictory to such a unified framework. However, volcanic rocksin the Yilgarn craton of similar age to those of the Superiorprovince were generally erupted through continental crustand, therefore, do not correspond to the more primitiveoceanic arc settings represented by the 2.7 Ga VMS-hostingterranes in Canada (Wyman et al., 1999).

In summary, the empirical association of mineral depositclasses with specific stages of the supercontinent cycle sup-ports the precept that mineral deposits are products of par-ticular geodynamic settings (Fig. 5).

Archean Geodynamics and Greenstone TerranesNeoarchean greenstone-granitoid terranes show both differ-

ences from and similarities to Proterozoic and PhanerozoicCordilleran-type orogenic belts that formed by terrane accre-tion at convergent margins (Burke et al., 1976; Sleep andWindley, 1982; Card and Ciesielski, 1986; Friend et al., 1988;Sengor, 1990; Sleep, 1992; Windley, 1995; Polat et al., 1999).Komatiitic liquids stem from melting in anomalously hot man-tle plumes. Their eruption temperature of 1,650ºC contrastswith ~1,200ºC for basalts. Komatiites are ubiquitous inArchean greenstone terranes but are rare in Proterozoic orPhanerozoic counterparts (Arndt, 1994). Together with basalts,they represent intraoceanic plateaus or continental floodbasalts. Given higher mantle temperatures in Archean plumes,plateau crust would have been thicker, ~30 to 50 km (Fig. 3)and thus not able to be subducted; rather, such crust was im-bricated where plateaus jammed against convergent margins(Bickle, 1986; Abbott et al., 1994a; Wyman et al., 1999).

At Archean convergent margins, bimodal arc magmatisminvolved slab dehydration and wedge melting, generating arcbasaltic liquids as in the Phanerozoic (Pearce and Peate,1995; Wyman, 2003). However, given their high thorium con-tents, trondhjemite-tonalite-granite (TTG) batholiths likelyformed as melts of enriched, garnet-amphibolite facies,plateau basalt crust subcreted beneath the convergent mar-gin, rather than depleted MORB-like crust (Foley et al.,2002). The TTG suite is characterized by a secular increase ofMg number and Ni from 4 to 2 Ga, conferring evidence of theinvolvement of a progressively thicker mantle wedge as sub-duction steepened (Martin and Moyen, 2002). Models of thethermal structure of the mantle predict a transition from flatto steep subduction at ~2.5 Ga, in keeping with the distribu-tion of TTG in Archean terranes and the transition in sedi-mentary rock REE patterns at this time (Abbott et al., 1994a;Taylor and McLennan, 1995). Given smaller plates, and acommensurately longer global ridge system in the Archean(Hargraves, 1986), ridge subduction would have been morefrequent, accounting for high heat flow in convergent mar-gins, which was responsible for the abundant TTG (Polat andKerrich, 2004).

Similarities between Neoarchean greenstone terranes andPhanerozoic convergent margins include accretionary tecton-ics, mélanges, subduction-accretion complexes, ophiolites,

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and Cenozoic-type arc associations. The Superior provincewas assembled by diachronous accretion in a Cordilleran-typeorogen from 2.74 to 2.65 Ga (Card and Ciesielski, 1986; Card,1990; Thurston et al., 1991; Percival et al., 1994; Calvert andLudden, 1999). A few small mélange occurrences have beendocumented in Archean terranes (Kusky, 1991; Wang et al.,1996; Polat and Kerrich, 1999), with mélanges indicating thepresence of subduction-accretion complexes. Precambrianophiolites, reviewed by Kusky (2004), indicate paleo-conver-gent margins. Boninites have been recorded from severalArchean volcanic rock sequences, as well as an association ofadakites, high Mg andesite, and Nb-enriched basalts, typicalof Cenozoic arcs that are linked to shallow subduction of rel-atively hot oceanic lithosphere (Kerrich et al., 1998; Hollings,2002; Polat et al., 2003).

Neoarchean greenstone belts are now generally consideredto be Cordilleran-style collages of oceanic arc and plateau ter-ranes, in which orogenesis was induced by plateaus jammingagainst arcs. The composite arc-plateau crust was stabilizedby the residue of plume melting, coupled to the compositecrust as continental lithospheric mantle (Wyman and Kerrich,2002). At Archean convergent margins, shallow subductionangles, ~11-km-thick oceanic crust (of which only the top ~7km was occasionally obducted) and relatively high thermalgradients, can explain the absence of blueschist-eclogite asso-ciations and rare ophiolites that generally lack a mantle sec-tion (Figs. 2E, 3; cf. Moores et al., 2000).

Metallogeny of Intraoceanic Arcs

Podiform Cr

Podiform bodies of spinel are an important resource ofchromium. Most of the deposits are in Caledonian or youngersuprasubduction zone ophiolites. Notable are the ~500, ~460,and ~370 Ma ophiolites of northwestern China, obductedduring accretion of arc terranes along composite sutures be-tween the Kazakhstan, Siberian, and Tarim blocks; Ap-palachian ophiolites; Hercynian ophiolites of Eurasia;Tethyan Mesozoic ophiolites, including those in Turkey,Oman, and Cyprus; and Mesozoic-Cenozoic ophiolites in ac-creted terranes of the North American Cordillera. Rare pod-iform chromitite bodies have been reported from a 3.0 Gaophiolite in the Ukraine, and the 2.5 Ga Zunhua ophiolite ofthe North China craton (Thayer, 1976; Duke, 1996a; Zhou etal., 2001; Polat et al., 2004).

Podiform bodies are dominated by Cr-rich spinels en-veloped by dunite in harzburgite of the mantle section, or thecrust-mantle transition, of oceanic lithosphere from intrao-ceanic arcs. Podiform morphology reflects mantle flow paths.A current model for development of chromitite bodies in-volves generation initially of hydrous basaltic melts in theperidotitic mantle wedge from dehydration of the subductingslab. Hydrous melts depolymerize, enhancing the octahedralsite preference for Cr3+

. Subsequent reaction of melt withperidotite in an open system induces polymerization accom-panied by precipitation of Cr spinel at ~7-km depth and 0.2GPa (Fig. 2C; Edwards et al., 2000).

Podiform chromite deposits reflect obduction of intrao-ceanic arc crust-upper mantle sections in both continent-con-tinent (Appalachian, Tethyan) and Cordilleran-type orogens

(Figs. 1, 2C). Sparsity of these deposits in Precambrian ter-ranes reflects the same process responsible for the absence ofblueschists and eclogites, or of complete ophiolite sections,given that the upper basaltic sections of thicker oceaniclithosphere were obducted (Fig. 2E; Moores, 2002; Polat etal., 2004).

VMS deposits

VMS deposits (Franklin et al., 2005) form in oceanicspreading centers, arcs, and rifts (Hannington et al., 2005),but mid-ocean-ridge crust is rarely preserved in the geologicrecord due to the likelihood that oceanic lithosphere will besubducted (Cloos, 1993). Many VMS deposits formed atconvergent margins under extensional conditions, specificallyin back arcs, where thinned and fractured lithosphere,upwelling asthenosphere, and high-temperature magmasgenerate long-lived high heat flow and enhanced hydraulicconductivity (Figs. 2C, 4E). Back-arc lithosphere is morereadily obductible, being young and hot. The fact that allVMS deposits are associated with some mafic magmatism sig-nifies a functional relationship to thermal anomalies in theupper mantle (Barrie and Hannington, 1999). A lack of sig-nificant VMS deposits in the Mesoproterozoic and Neopro-terozoic (Hutchinson, 1981; Meyer, 1981, 1988) reflects thedrift stage in dispersal of first Columbia and then Grenvilleorogens that stitched together Rodinia. These orogens nowexpose deep erosional levels, which is ultimately due to de-lamination of mantle lithosphere (Fig. 5).

Based on rock associations, and therefore tectonic setting,Barrie and Hannington (1999) and Franklin et al. (2005)classified VMS deposits into five groups. Mafic and bimodalsiliciclastic rock-associated deposits are mainly restricted tothe Phanerozoic. The former consists of tholeiitic with minorboninitic rocks and includes ocean-ridge deposits that wereobducted as part of ophiolite fragments, exemplified byTethyan ores of Cyprus and Turkey. The geodynamic settingis a suprasubduction zone, and such magma-ore associationsextend to the Paleoproterozoic Flin Flon VMS province(Wyman, 1999). The latter, characterized by large tonnageswith high Pb but low Cu contents, formed in a continentalarc or back-arc setting; VMS ores of the Bathurst and Iber-ian Pyrite Belt provinces are prominent examples of thisgroup.

The other three groups of VMS deposits have broader sec-ular distributions. Bimodal-mafic and bimodal-felsic groupdeposits occur in oceanic terranes back to the Neoarchean ofsome cratons. The former represent primitive oceanic arcs orback arcs; examples include Noranda and Matagami, Quebec,some ores of Flin Flon, Saskatchewan, and Manitoba, andJerome, Arizona. The latter represents precipitation of VMSdeposits in mature arcs, such as the Mt. Read district, Tasma-nia. A mafic volcanic-volcaniclastic rock and turbidite associ-ation with VMS formation occurred from the Mesoprotero-zoic through the Phanerozoic. These deposits developed insediment-rich oceanic rifts, notably Windy Craggy, BritishColumbia, or in propagating continental rifts, exemplified bythe Besshi district of Japan. The Middle Valley and Escanabatrough, and the Sea of Cortez, are present-day metal-richanalogs to these two environments in the final group, respec-tively (Barrie and Hannington, 1999).

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Intraoceanic and Continental Margin Arc Porphyry-Epithermal Systems

Porphyry Cu-Mo-Au (hereafter referred to as porphyry Cu)and related epithermal Au-Ag deposits are predominantly, butnot exclusively, a Phanerozoic occurrence (Seedorff et al.,2005; Simmons et al., 2005). The majority of both deposit typesoccur in Mesozoic and Cenozoic subduction-related subvol-canic plutonic complexes and related volcano-sedimentary se-quences, but this may be in part a function of the low preser-vation potential of shallow-level crustal sequences within activeconvergent plate margins. Rapid uplift and erosion, tectonicerosion, and collision (either with oceanic terranes such as is-land arcs, seamounts, or plateaus, or with continental masses)commonly result in destruction of supracrustal sequences inboth oceanic and continental volcanic arcs. Nevertheless, de-posits of both types do occur in older terranes, but with in-creasing rarity back to the Mesoarchean, to the point that Pre-cambrian occurrences in Australia, Canada, India, andScandinavia are noted as exceptions; the earliest known de-posits are ~3.3 Ga in age (Barley, 1982). The characteristics ofPrecambrian deposits are little different from those of theirPhanerozoic counterparts (Gaál and Isohanni, 1979; Barley,1982; Roth et al., 1991; Fraser, 1993; Sikka and Nehru, 1997;Stein et al., 2004), suggesting that similar tectonomagmaticprocesses were involved in their formation.

Porphyry Cu deposits

Porphyry Cu deposits show one of the clearest relationshipsto specific plate tectonic processes of any ore deposit type(Fig. 6; Sillitoe, 1972; Burnham, 1981). The relationship tosubduction of oceanic crust relates primarily to the large fluxof water and other volatiles from the slab into the overlyingasthenospheric mantle wedge. As recently reviewed byRichards (2003; see also Candela and Piccoli, 2005), thesevolatiles metasomatize the mantle wedge and reduce its melt-ing point, such that hydrous basaltic magmas are produced bypartial melting in the highest temperature regions. Thesemelts are the ultimate sources of more evolved magmas thatare emplaced into the overlying crust and which may gener-ate porphyry and related epithermal deposits.

Subduction represents the return flow of materials into themantle to compensate for the creation of new oceanic lithos-phere at mid-ocean ridges. But processes of sea-floor meta-morphism, resulting in hydration and introduction of othersea water-derived elements, such as S, Cl, and alkalis (ex-changed for Ca), mean that the return flow is modified fromthe original MORB composition. Upon return into the man-tle, these same water-soluble elements are released duringprograde dehydration reactions, whereby minerals such asserpentine, amphibole, chlorite, zoisite, and lawsonite (Fig.6;Tatsumi, 1986; Schmidt and Poli, 1998; Winter, 2001;Forneris and Holloway, 2003) are converted to progressivelymore anhydrous blueschist- and eclogite-facies assemblages.Additional components may be added by subduction of sea-floor sediment and tectonic erosion of upper plate rocks (e.g.,de Hoog et al., 2001).

Basaltic crust of the downgoing slab may partially meltwhere the lithosphere is young (<25 m.y.; Defant and Drum-mond, 1990; Peacock et al., 1994) and/or during shallow

subduction, both of which may result in higher temperaturesbeing achieved in the slab at shallow depths. Normal subduc-tion of oceanic lithosphere results predominantly in dehydra-tion and release of a water-rich fluid phase into the overlyingmantle wedge. Fluid release probably begins at the shallow-est levels of subduction but appears to reach a maximum atdepths of ~100 km, corresponding to the final breakdownconditions of serpentine, amphibole, and chlorite, all of whichappear to have maximum stabilities at ~3 GPa and 700° to850°C (Schmidt and Poli, 1998). These depths correspond tothe characteristic depth of the Benioff zone beneath volcanicarcs, suggesting a direct connection between slab dehydrationand magma generation. Micas may persist to greater depthsand higher temperatures, which may, in part, explain the ob-served K2O increase in magmas toward the back arc (Schmidtet al., 2004).

Convection of metasomatized peridotite into warmer cen-tral parts of the mantle wedge, or direct fluid infiltration, re-sult in partial melting to form high Mg basalts with as muchas 2.5 wt percent H2O, enrichments in large ion lithophile el-ements, relatively high oxidation state compared with MORB(as much as two log units above fayalite-magnetite-quartz),and high sulfur contents (experiments suggest S concentra-tions as high as ~1.5 wt % in oxidized basaltic melts; Jugo etal., 2005).

Concentrations of chalcophile and highly siderophile ele-ments in these primary melts may be controlled by the stabil-ity and abundance of residual sulfide phases in the mantlewedge source, which is in turn a function of oxidation state(Candela, 1992). With increasing oxidation state, concentra-tions of chalcophile elements, such as Cu, will reach a maxi-mum prior to the concentration of highly siderophile ele-ments, such as Au and platinum group elements (Richards,2005, and references therein). This observation may explainsome of the variation in Cu/Au ratios in porphyry systems,with Cu-rich deposits being generated under normal subduc-tion conditions leading to moderate mantle wedge oxidation,and Au-rich deposits being formed under more extreme oratypical conditions that result in complete destruction ofresidual sulfide phases either by extreme oxidation or multi-ple stages of partial melting (e.g., during tectonic transitionssuch as subduction-polarity reversal or termination, back-arcextension, or arc collision; Solomon, 1990; Wyborn and Sun,1994; Richards, 1995). Mungall (2002) has recently suggestedthat highly oxidized adakite magmas produced by slab melt-ing may also have this effect of sulfide destruction.

Processes affecting the composition of primary subduction-related magmas thus appear to be the most fundamental con-trols on metallogenesis in volcanic-plutonic arcs. Althoughnot all arcs or magmatic suites within arcs host economic por-phyry Cu deposits, few deposits are known that cannot beclearly related to subduction magmatism or to magmas de-rived from subduction-modified mantle; a possible exception,which is associated with a continental rift, is described byBlecha (1974). Because the exsolution of metalliferous hy-drothermal fluids occurs in the final stages of magmatic evo-lution, many factors can intervene between initial magmageneration and upper crustal emplacement to affect the ore-forming potential of these magmas and their exsolved fluids.The simplest constraint is the magmatic flux into the upper

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crust. If the rate and volume of supply of magma is limited,then so too will be the flux of heat, metals, and other ore-forming components (Fig. 6). This constraint implies that thelargest porphyry systems will be associated with long-livedand voluminous arc magmatic suites.

Tosdal and Richards (2001) and Richards (2003) reviewedstructural controls on the emplacement of porphyry magmasin the upper crust and argued that tectonic stresses acting on

a regional architecture of translithospheric structures may in-fluence the location of magma ascent by providing relativelypermeable pathways. Optimal sites are extensional structuraldomains formed at jogs and stepovers in large strike-slip faultsystems deforming under mildly oblique compressional stress(Fig. 6B). Although magma ascent can occur in the absenceof such structures, their existence may act to focus magmaflux, thus enhancing subsequent ore-forming potential. A

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FIG. 6. A. Normal subduction configuration beneath a continental arc (from Richards, 2003; modified from Winter, 2001).Slab dehydration leads to hydration of the overlying asthenospheric mantle wedge and partial melting in the hotter centralregions of the wedge. Hydrous basaltic melts pool at the base of the crust due to density contrasts, where they fractionate,release heat, and interact with crustal materials to generate more evolved, less dense andesitic magmas (by melting, assimi-lation, storage, and homogenization—MASH process of Hildreth and Moorbath, 1988), which can then rise to upper crustallevels. It is these evolved magmas that are directly associated with porphyry Cu deposit formation. B. Oblique convergenceleads to the generation of structurally permeable transpressional sites along trench-linked strike-slip faults, up which magmamay ascend from lower crustal MASH zones. Rapid, voluminous emplacement of magmas in the upper crust is regarded hereto be a prerequisite for the subsequent formation of large porphyry Cu deposits by magmatic-hydrothermal fluid exsolution.

spatial relationship of ore deposits to such structural nodes,often recognized in regional exploration as lineament inter-sections, has been noted in many porphyry and related ep-ithermal districts (e.g., Corbett and Leach, 1998; Sasso andClark, 1998; Padilla Garza et al., 2001; Richards et al., 2001;Chernicoff et al., 2002; Sapiie and Cloos, 2004).

Other models for porphyry Cu formation have invoked thedirect involvement of slab melts (adakites; Sajona and Maury,1998; Oyarzun et al., 2001) or the role of crustal thickeningand shallowing of subduction angle in affecting magma gen-eration and composition (Kay et al., 1999). However, al-though these processes may be important locally, they do notseem to be universally applicable, and a more general rela-tionship to subduction magmatism is implied. Variationsamong the porphyry suite may arise from the wide variety ofpossible tectonic configurations in subduction zones, and spe-cific events or combinations of events may cumulatively act tomaximize (or reduce) porphyry-forming potential. Notably,differences in porphyry systems in oceanic versus continentalarcs occur mainly in subtle details and not in overallprocesses. Oceanic arc systems tend to be associated withsomewhat more mafic (dioritic) plutonic rocks, whereas con-tinental arc systems are typically associated with more felsicsystems (Hollister, 1975; Kesler et al., 1975). There is a com-mon tendency for oceanic systems also to be somewhat moreAu versus Mo rich in continental systems, although many ex-ceptions exist. Both of these variations may relate to the de-gree of fractionation and crustal interaction experienced bythe primary magmas (oceanic systems representing moreprimitive systems) and continental porphyries being morefractionated (loss of Au) and contaminated with crustal com-ponents (higher Mo; Farmer and DePaolo, 1984; Blevin andChappell, 1992).

Epithermal Au-Ag deposits

Historically, an understanding of the relationship betweenshallow-level epithermal Au-Ag deposits and subvolcanic por-phyry systems was slower to develop than the overall rela-tionship to convergent plate margins. This was primarily dueto problems of preservation and exposure level, which meantthat where near-surface deposits were preserved, erosion hadnot penetrated deeply enough to reveal underlying mag-matic-hydrothermal systems. Conversely, where porphyry de-posits were exposed, overlying epithermal deposits had al-ready been removed. Consequently, near-surface advancedargillic alteration, characteristic of high-sulfidation–type ep-ithermal deposits, was not included in the classic model ofporphyry alteration and mineralization zoning of Lowell andGuilbert (1970). Nevertheless, Sillitoe (1973) made an earlyconnection between porphyry formation and surficial vol-canic and fumarolic activity, and later studies, such as those ofthe adjacent Far Southeast (porphyry) and Lepanto (high-sul-fidation epithermal) deposits by Arribas et al. (1995) andHedenquist et al. (1998), clearly demonstrated a connectionbetween these distinct ore-forming environments. As such,the tectonic controls on high-sulfidation epithermal mineral-ization are closely related to those affecting porphyry de-posits. However, economic deposits of both types need notform together, because local details of fluid evolution, trans-port, and deposition processes may favor ore deposition in

one or the other environment but not necessarily both envi-ronments. For example, Bissig et al. (2002) recently proposedthat regional uplift and erosion history was critical in control-ling the development of mineralized epithermal systems inthe El Indio-Pascua belt (Chile and Argentina), which are as-sociated only with apparently barren plutons. Thus, drillingbeneath a known epithermal deposit will not necessarily re-veal an economic porphyry deposit, although evidence of ahigh-temperature magmatic hydrothermal system is likely tobe encountered.

Unlike high-sulfidation systems, low-sulfidation epithermaldeposits do not show a clear, exclusive relationship to sub-duction zone magmatism, and many deposits are generatedby thermal anomalies caused by crustal extension, such asepithermal Au-Ag deposits in the Basin and Range district,Nevada (Berger and Bonham, 1990; John, 2001; Simmonset al., 2005). In this respect, the involvement of specificmagmatic components (both volatiles and metals) in low-sul-fidation epithermal systems is less clear, and the key input forsuch systems may simply be a heat source of any origin. Bycontrast, intermediate-sulfidation epithermal systems arecommonly found in porphyry districts, and either a direct ordistal association with magmatism has been proposed in manyinstances (e.g., Rye, 1993; Hedenquist et al., 1996; Hayba,1997; Faure et al., 2002). A common structural control onmost epithermal-type deposits is extensional faulting andbrecciation, either generated regionally by tectonic stressfields (as in the case of the Basin and Range) or locally byforces involved with magma emplacement (crustal doming)or by elevated fluid pressure (hydraulic fracturing). The lattertectonic condition is commonly generated in association withporphyry formation but not exclusively so.

Metallogeny of Cordilleran Orogens

Metallogenic context

In contrast to the shallow crustal regions that characterizecontinental magmatic arcs, as described above, much of anevolved orogen exposes rocks that were deformed and meta-morphosed at deeper crustal levels. Crustal rocks that wouldhave hosted porphyry and related epithermal mineral de-posits are typically unroofed and eroded in fore- and back-arcregions. The exposed middle crustal rocks in these regions aredominated, in contrast, by mineral deposits that reflectdeeper hydrothermal processes that are active in convergentto transform continental margins. These processes formmainly orogenic Au deposits, with commonly related As, W,Sb, and Hg resources. In addition, preaccretionary mineraldeposits, such as podiform Cr and VMS deposits that weredescribed above, may also be present and hosted within thesame blocks of accreted juvenile crust (Fig. 1A).

High heat flow and intense fluid regimes are important tec-tonic features inherent to most Cordilleran orogens. The gen-eration of Barrovian P-T conditions is typical for progressiveaccretion of a broad zone of radiogenic juvenile materialscraped off a downgoing slab, where clockwise P-T-time tra-jectories generate deeper and later metamorphism. Underthese heat-flow conditions, peak metamorphism at mid-crustallevels (greenschist facies) predates peak metamorphism inthe deeper crust, such that fluids generated by dehydration

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reactions in deeper crust advect to the mid-crust where theyoverprint the peak-metamorphic assemblage (McCuaig andKerrich, 1998). Within about 15 m.y. of accretion, large areasof the mid-crust will begin to experience a significant rise ingeotherms (e.g., Jamieson et al., 1998). The causes of thethermal episode are complex; increased radioactive heat pro-duction of accreted material is the most commonly cited trig-ger, but shear heating, massive fluid flow, crustal thickening,ridge subduction, or slab rollback are all processes that mayadd heat into the growing continental margin. Rapid upliftand/or continued outboard subduction typically yields a pat-tern of inverted isotherms, such that more highly metamor-phosed rocks are thrust above lower grade rocks (Peacock,1987).

Fluid reservoirs are present both in the subducted slab, asnoted above, and in the accreted sedimentary and volcanicrock sequences. As described above, slab devolatilizationreleases volatiles into the overlying mantle wedge. The pro-grading accreted juvenile crust represents a significant sec-ond reservoir, with voluminous fluid release across variousmetamorphic isograds (Fyfe et al., 1978; Powell et al., 1991).Estimates for progressive metamorphism of an average peliteare that about 5 vol percent of the rock will be lost to the fluidphase at metamorphic reaction boundaries (e.g., Walther andOrville, 1982). Fluids released at greenschist- and amphibo-lite-facies conditions typically consist of H2O, CO2, CH4, andN2 (Mullis, 1979), as well as relative enrichments of H2S fromdesulfidation reactions (Ferry, 1981), therefore explaining thedominance of C-O-H-N-S fluids in Cordilleran orogens. Spe-cific volatile composition of these fluids generated duringmetamorphism will depend on the composition of the juve-nile rocks, particularly on the clay, carbonate, and organicmatter content (Yardley, 1997). Numerous studies (see sum-mary by Goldfarb et al., 2005) also indicate a progressive mo-bilization of As, Au, B, Hg, Sb, and W in such fluids with in-creasing degree of metamorphism. Concentration of thesespecies in metamorphic fluids may determine, to a large part,

mineral resource potential within Cordilleran orogens. Silicametasomatism in both the mantle wedge and overlying crustis commonplace (Manning, 1997) and, as a result, there is aconsistent association of epigenetic ore deposits in metamor-phic environments with large quartz vein systems (Fig.7).

Orogenic Au

The fore-arc regions of Cordilleran orogens inherently arecharacterized by widespread orogenic gold deposits. The typeCordilleran orogen of western North America, which is stillevolving, may have begun to form anywhere from 400 to 200m.y. ago, depending on how an orogen is defined. Subsequentto Rodinian rifting in the Neoproterozoic, the Pacific marginof North America was the passive margin site of sedimenta-tion through the Middle Devonian (Dickinson, 2004). By theLate Devonian, convergent tectonism began along the mar-gin, with the Late Devonian to Early Mississippian Antler (orEllesmerian in the far north) and Late Permian to Early Tri-assic Sonoma allochthons of oceanic rocks being thrust overthe miogeoclinal shelf edge (Burchfiel et al., 1992). Such ob-duction of oceanic rocks was not associated with any type ofsubduction zone geodynamics, continental arc development,or metamorphism, and this low-temperature tectonism alsolacked any associated ore deposit formation of significance.Cordilleran orogenesis essentially began with the accretion ofmore than 200 terranes along the seaward side of the formerpassive margin post-Early Triassic (Fig.1A; Coney et al., 1980;Monger et al., 1982). The exact time of initiation of simulta-neous slab subduction and terrane accretion, and thus thebest estimate of the start of orogenesis, could be any time be-tween ~240 and 70 Ma. Moores et al. (1999) noted that thereis a lack of evidence for such terrane collision along much ofthe margin prior to the younger part of this age range.

With the onset of subduction-accretion and the deeper andlater style of metamorphism, economically significant oro-genic Au deposits have formed within mainly greenschist fa-cies rocks of the Cordilleran orogen for probably the last 170

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FIG. 7. Cordilleran-type orogens are recognized for the widespread distribution of orogenic gold deposits in metamor-phosed juvenile rocks on either side of the magmatic arc. Ore-forming fluids in the fore arc may be derived from progrademetamorphism of accreted material above a subducting slab and from the slab itself; where slab fluids are released into themantle wedge, mantle-derived melts may carry some of the fluid into the accreted oceanic rocks. The metalliferous fluidsare focused along major crustal shear zones in the fore arc, which previously may have been sites of terrane suturing.

m.y. (Fig. 5; Goldfarb et al., 2001). The youngest such oresare the ~50 Ma gold deposits on Chichagof Island, southeast-ern Alaska. However, it is likely that younger orogenic golddeposits have formed at depth within the fore arc of the oro-gen since the middle Eocene, but these mid-crustal ore-host-ing regimes have yet to be uplifted and exposed at the surface(Fig. 7; Goldfarb et al., 2000).

The most significant lode deposits are associated with ter-rane-bounding fault systems. Where no such major conduitsoccur within a deeper and later thermal sequence, veins aresmaller and more widely distributed, and world-class economicgold lodes are unlikely to have formed (e.g., Chugach Moun-tains/Kenai peninsula, Nome, Klondike). Both exotic oceanicblocks, such as hosts for the Mother Lode and Juneau goldbelt, and terranes of pericratonic miogeoclinal strata, includingthe Fairbanks and Klondike districts in the Yukon-Tanana ter-rane, all of which were translated along the North Americanmargin, are equally likely to host orogenic gold deposits.

The oldest gold lodes in the North American Cordillera arethose of Middle and Late Jurassic in the Canadian sector andLate Jurassic to Early Cretaceous in California (Goldfarb etal., 1998). In Alaska, both gold and arcs young seaward, from~100 Ma in the north and interior to ~60 to 50 Ma along thepresent-day active margin. Typically, the orogenic goldprovinces occur at geologic and structurally favorable loca-tions in terranes of the fore arc, such as within the Juneaugold belt and Sierra foothills. However, where arcs are rela-tively diffuse, rather than occurring as distinct Andean-stylebatholiths, important lodes occur within an evolving arc (in-cluding hosts to deposits of the Klamath Mountains and Fair-banks districts). Where well-defined batholiths have alreadybeen crystallized and are in the process of regional uplift,competent margins to these igneous masses may also hostorogenic gold deposits (e.g., Willow Creek). In addition to anumber of small orogenic gold deposits in the Cordilleranback-arc regions (e.g., Polaris-Taku, northern British Colum-bia; Humboldt Range, Nevada), the world-class Late Creta-ceous Bridge River deposit in southern British Columbia in-dicates important orogenic gold ore formation, as well assubduction-related plutonism, may also continue landwardinto oceanic terranes inboard of an evolving continental mar-gin arc. The thermal profile of a Cordilleran orogen, ratherthan simply a geographic location in a growing margin, ap-parently controls fluid evolution and ore genesis in theoceanic rocks (McCuaig and Kerrich, 1998). Indeed, a similararc to back-arc position characterizes many of the Late Juras-sic-Cretaceous orogenic gold deposits in the deformed terri-geneous rocks to the west of the Siberian craton in easternRussia (Fridovsky and Prokopiev, 2002).

The Altaid orogen presents a similar Au-rich Cordilleran-type orogen composed of Vendian through Jurassic unitsaccreted to the margins of the Siberian craton (Sengor andNatal’in, 1996b). Inclusion of the Baikalides and Uralides,both containing important Paleozoic orogenic gold provinces,remains controversial (Sengor, 1993). Tectonism and defor-mation span the entire duration of the Paleozoic. Giant EarlyPermian orogenic gold deposits (e.g., Muruntau, Zarmitan,Kumtor, Sawyaerdun) continue along the length of the oro-gen in what is probably one of the outermost accreted ter-ranes (Yakubchuk et al., 2002). In a pattern similar to that

observed in Alaska, older orogenic gold provinces reflect ear-lier subduction closer to the craton margin. Giant depositssuch as Olympiada and Zun-Kholba formed in Proterozoicterranes along the southwestern side of the craton in the lat-est Neoproterozoic and early Paleozoic, followed by ores inmore seaward regions of Kazakhstan and the Urals in themid-Paleozoic, and then the Permian ores developed alongthe edge of the closing Paleo-Tethyan Ocean (Herrington etal., 2005; Yakubchuk et al., 2005).

Significant characteristics of the Altaid orogen (Yakubchuket al., 2005) illustrate other broad tectonic controls on oro-genic gold in Cordilleran orogens. First, the immense gold re-source at the Sukhoi Log deposit, probably of mid-Paleozoicage (Goldfarb et al., 2001), is hosted by carbonaceous andpyrite-rich flysch in a retroarc location within complexly de-formed Neoproterozoic pericratonic Baikal terranes (Bulga-tov and Gordiyenko, 1999). The thermal event associatedwith emplacement of the immense Angara-Vitim batholith(Yarmolyuk et al., 1998) correlates with the major period oforogenic gold deposit formation within 100 km of the craton.Thus, there are clearly significant exceptions to the generalobservation that orogenic gold ores in a Cordilleran orogenwill always be younger in an oceanward direction. Second,with the exception of this Baikal region, large gold placers,such as those that dominate the circum-Pacific goldfields, areabsent. Perhaps this reflects the fact that continent-continentcollision closed the Altaid orogen and has, at least temporar-ily, formed a Paleozoic craton. This preserved paleo-Cordilleran margin has thus not been susceptible to rework-ing and erosion of significant amounts of its contained lodegold systems. Further support for such a concept is that muchof the interior of the Altaid orogen still contains numerousPaleozoic porphyry and epithermal deposits (Yakubchuk etal., 2002, 2005), whereas such shallow crustal levels havebeen already removed by uplift and erosion from many of thecircum-Pacific Cordilleran terranes.

The Paleozoic Tasman orogen of eastern Australia, whichincludes the gold-rich Thomson, Hodgkinson-Broken River,and, particularly, Lachlan fold belts, may also be consideredan accretionary orogen but with important differences fromthe more classic Cordilleran-type orogens of western NorthAmerica and the Altaids. Rather than a series of accreted ter-ranes, much of the more deformed and metamorphosed sec-tors of the orogen reflect a single, quartz-rich turbidite fansystem shed off the Delamerian-Ross highlands in the earliestPaleozoic. Ordovician-Silurian orogenesis was dominated byshortening and folding, as is typical of Cordilleran orogens,but these were thin-skinned tectonic events and lacked anymajor uplift of basement blocks (Coney, 1992; Goldfarb et al.,1998). This difference in crustal response may be indicativeof subduction and/or accretion in association with a large fansystem, rather than a series of terranes, along a continentalmargin (Gray and Foster, 2000). The extensive ores of theVictorian goldfields formed during Late Ordovician deforma-tion, metamorphism, and subduction in the western provinceof the Lachlan fold belt (~440 Ma: Bierlein et al., 2001); how-ever, no magmatic arc developed during subduction beneaththe deforming turbidite wedge (Fergusson, 2003).

Thrust-fault development and uplift of the Victorian orehost rocks began at ~455 Ma, with perhaps slab rollback ~15

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m.y. later, providing the main thermal event related to goldformation (Squire and Miller, 2003). Therefore, the Tasmanorogen scenario suggests that deformation, heating, and upliftof juvenile material along a margin may be essential to fluidproduction, fluid migration, and related lode gold formation,regardless of the presence of associated magmatism or anabundance of well-defined terrane-bounding fault zones.

The Otago schist belt of the South Island of New Zealandmay be more like a classic Cordilleran orogen, or at least apart of such an orogen, but also without any magmatic arc ac-tivity recognized in the gold-hosting terranes. This Permian-Cretaceous accretionary wedge contains a number of terranesthat likely amalgamated and were simultaneously deformedand metamorphosed in Late Jurassic-Early Cretaceous (Mor-timer, 1993; Gray and Foster, 2004). Orogenic gold formationoccurred in the schists during this deformation and uplift,probably at ~150 to 130 Ma (Craw, 2002). This time is ap-proximately mid-way through the 150-m.y.-long episode ofterrane translation along the margin of East Gondwana(Pickard et al., 2000), such that hydrothermal activity oc-curred within the actively deforming rocks as they were partlybetween their original location off the northeastern coast ofAustralia and present South Island location.

Another variation on a Cordilleran-style margin might beeast-central Asia, where terranes that now form the Japaneseislands and southeastern Russia were at one time immediatelyseaward of the eastern margin of China and have since un-dergone significant Jurassic(?)-Cretacous strike-slip transla-tion (Sengor and Natal’in, 1996b; Charvet et al., 1999). Theresulting Cenozoic configuration, subsequent to northwardmigration of the entire subduction and/or accretion complex,includes rocks of the North China craton now located imme-diately along the Pacific margin. These Precambrian rocks, aswell as the migrated Mesozoic sequences, contain importantorogenic gold deposits; indeed, the orogenic gold deposits inthe North China craton represent the only known significantPhanerozoic gold ores in any Precambrian craton (Goldfarbet al., 2001; Zhou et al., 2002). Gold ores along the northern,eastern (i.e., Jiaodong), and southern (i.e., Qinling) marginsof the craton formed at ~130 to 120 Ma, during delaminationof the eastern half of the Archean continental lithosphericmantle (Griffin et al., 1998). Delamination of continentallithospheric mantle, mantle magmatism (generally at ca.160–125 Ma), and hydrothermal activity may relate to slabsubduction from the north and south, and/or circum-Pacificoblique subduction along the eastern transform margin, allduring the Mesozoic. Because these rocks were highly meta-morphosed already 2 b.y. prior to this Yanshanian orogen,ore-forming fluids must have been sourced in either the un-derplated material or the mantle melts, although the gold it-self still could have a crustal source. The North China exam-ple indicates that decratonization, or delamination, ofPrecambrian terranes during some form of continental mar-gin tectonism may still lead to formation of orogenic gold sys-tems, despite the presence of crustal rocks that were firsthighly tectonized and devolatilized billions of years earlier.

Cordilleran-style orogens of Precambrian age, althoughmore difficult to recognize, also inherently contain wide-spread orogenic gold deposits in, most commonly, green-schist-facies terranes. Neoproterozoic gold-forming events,

such as the Pan-African of the Nubian-Arabian Shield, showlittle difference from younger Phanerozoic accretionary oro-genesis. Subduction-accretion in eastern Africa occurredfrom 900 to 690 Ma, with the orogen dominated by the addi-tion of new juvenile material (approx 80%), and this terraneaccretion was then followed by 100 m.y. of strike-slip onmajor fault systems (e.g., Stern, 1994; Genna et al., 2002),which corresponds to the onset of oblique convergence af-fecting much of Gondwana (Veevers, 2003). Widespread golddeposits formed throughout the shield at this time (Albino etal., 1995; LeAnderson et al., 1995), which probably includedthe great goldfields of ancient Egypt-Nubia (Klemm et al.,2001), and other important ores along the southern margin ofthe opening Tethyan ocean basin such as the Tuareg andNigerian Shields of West Africa. The Halls Creek orogen ofnorthern Australia, host to the Telfer deposit, is also of thisage.

The same style of accretionary tectonics appears to havecharacterized the Paleoproterozoic, as in the 2.1 Ga Birimianterrane, West African Shield; 1.8 Ga Homestake deposit,South Dakota, of the Trans Hudson orogen; and 1.7 Ga Ash-burton and Pine Creek gold provinces of northern Australia(Hirdes et al., 1996; Attoh and Ekwueme, 1997; Sener et al.,2005). Similarly, gold-rich provinces in the broad,Neoarchean, Cordilleran-type superfamily of accretionaryorogens have been documented in the Superior province(Kerrich and Wyman, 1990; Polat and Kerrich, 2001), Yilgarncraton (Myers, 1993, 1995), Slave province (Kusky, 1989), andZimbabwe cratons (Kusky, 1998), where major terrane-bounding faults focused auriferous fluids (Kerrich and Feng,1992; de Ronde et al., 1997).

The span of Earth history from ~1.8 to ~0.8 to 0.6 Ga is no-table for the lack of orogenic gold deposits, despite numerousCordilleran orogens, particularly as products of the Mesopro-terozoic growth of Rodinia. However, due to the cessation ofextensive cratonization (Archean-type continental lithos-pheric mantle) on a cooling Earth, much of these continentalmargin orogens were not preserved and only high-grade base-ment rocks remain (Goldfarb et al., 2001). These sectors ofthe Cordilleran orogens lie beneath typical gold-favorable en-vironments, and thus more than 1 b.y. of orogenic gold for-mation has apparently been eroded from the geologic record.In addition, some of the Rodinian margins were composed ofterranes of mainly continental affinity, exemplified by theGrenville province (Condie and Chomiak, 1996), whichwould have lacked the volatiles, and perhaps gold, whichoccur in more juvenile oceanic terranes and are critical forore formation.

W and As

Scheelite is almost universally associated with orogenicgold provinces (Boyle, 1979) in more deeply unroofed meta-morphic environments. Many of the largest orogenic gold de-posits (e.g., Muruntau, Hollinger-McIntyre, Olympiada, Mt.Charlotte) contain notable W enrichments, with scheelite sig-natures commonly being relatively HREE and Sr rich, andMo poor (Kempe and Oberthur, 1997). The As-Au-B-Sb-Wgeochemical signature, with W at economic concentrations inplaces such as Yellow Pine, Idaho (Cookro et al., 1988), char-acterizes epigenetic ores in most orogenic belts. Historically,

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minor amounts of W have been produced from some shearzone-hosted quartz veins in orogenic gold provinces, includ-ing the Sierra foothills, California, and Otago schists, SouthIsland, New Zealand (Henley et al., 1976). In the Variscan ofcentral and southern Europe, W-rich ores have been particu-larly economically significant at Mittersill, Austria, the Bo-hemian Massif, Czech Republic, and Chataigneraie, France.

Recent data indicate that these are epigenetic deposits,which formed during deeper and later metamorphic events,surplanting the syngenetic model of some workers (Marignacand Cuney, 1999). The spatial association between gold andscheelite in Phanerozoic ore systems is further supported bythe consistent association of abundant scheelite grains inmany placer goldfields.

The spatial association between W and Au in metamorphicbelts, which has been well recognized for decades (Boyle,1979), likely reflects similar solubilites of the two trace ele-ments in low-salinity, aqueous-carbonic crustal fluids (Foster,1977). However, in Precambrian orogens, although green-stone belts can be sources for gold ores, most associated maficto ultramafic lithologic units would have relatively low con-centrations of W and are unlikely sources for W enrichmentsin hydrothermal fluids. Thus, hydrothermal systems that haveinteracted with basement granitoids may be critical for ex-plaining tungsten-rich quartz vein systems in greenstone belts(e.g., Zimbabwe craton: Foster, 1977).

Arsenic resources are associated with both a variety of mag-matic arc-related, shallow-formed mineral deposits (de-scribed in previous sections) and with auriferous veins inmetamorphosed terranes landward and seaward of the arc.Currently, much of the world’s As resources are in differentdeposit types in China; however, historically, many Precam-brian and Phanerozoic shear zone-hosted orogenic gold de-posits were the source for As (Fig. 7).

Hg and Sb

In areas of limited erosion within accretionary orogens, Hg,Hg-Sb, and Sb deposits, typically with anomalous gold, arepreserved as the most characteristic resources formed withinthe upper few kilometers of the allochthonous terranes(Studmeister, 1984; Dill, 1998). Such deposits, within low-grade to unmetamorphosed host units, have precipitatedfrom the same CO2-, 18O-, N-, and Cs-, Rb-, and K-rich fluidtypes that likely formed the Au, W, and As concentrations inthe higher P-T environments (Fig. 7; Goldfarb et al., 1990,2005). Seismicity along active continental margin faults leadsto discharge of such deep crustal fluids at the Earth’s surface,as along the San Andreas and Alpine, New Zealand fault sys-tems (White, 1967, 1981). If associated hydrothermal systemsare large and far travelled, then these Hg- and Sb-rich sys-tems may be important indicators of deeper orogenic Au de-posits. Such metallogenic zoning of Cordilleran-style veinsystems has been suggested by many workers (Kerrich, 1987;Nesbitt and Muehlenbachs, 1989; Ashley and Craw, 2004).The consistent Hg-Sb-As-Au association within such orogensreflects a similar affinity to bisulfide complexing. Lindgren(1895) pointed out more than 100 years ago that features suchas the trace element suites, quartz-carbonate-pyrite gangue,and carbonate-dominant alteration were indicators that oro-genic gold deposits in the Mother Lode province and epi-

zonal mercury deposits in the California Coast Ranges hadsome type of important genetic association.

Given their shallow level of formation, Hg ores are poorlypreserved. With the exception of the giant 400 Ma Almadendeposit in Spain, other globally productive mercury systemsare ~220 Ma or younger (Obolenskiy and Naumov, 2003).Genesis of the Almaden Hg province remains problematic:sea-floor exhalative, mantle plume, and epigenetic shearzone-related tectonism have all been proposed (Hernandez etal., 1999). However, the structural setting, in conjunction withCr mica-altered peridotite, is consistent with a Cordilleran-type suture (Jebrak et al., 2002). Mercury deposits of thePermo-Triassic of central Asia, Jurassic-Early Cretaceous ofeastern Russia, Late Cretaceous of southwestern Alaska, andCenozoic of the western United States are less problematic;all appear to be products of tectonism in evolving orogens.The association of Hg ores at Almaden and in the Dometskbasin anticline in Ukraine, Europe’s second most historicallyproductive Hg province, with older, transcrustal fault systems(de Boorder et al., 1995), is consistent with Caledonian andVariscan age orogeny, respectively (Fig. 7).

There is an abundance of Sb deposits, with notably few spa-tially associated Hg-rich systems, throughout the Acadian-Hercynian domain of the Canadian Appalachians and muchof western and central Europe (Mossman et al., 1991). Thisrelationship suggests catchment erosion of the shallowest partsof the Gondwanan host terranes, such that few near-surfaceores have been preserved beyond supercontinent breakup.

Placer Au

Economic concentrations of Cenozoic placer gold charac-terize catchments of both the fore- and back-arc regions oforogens throughout the circum-Pacific (Henley and Adams,1979; Goldfarb et al., 1998). These include the great gold-fields of the Mother Lode (Calfornia), Klondike (Yukon),Fairbanks and Nome (Alaska), the Russian Far East, andOtago (New Zealand). The source of gold is in auriferousveins of Jurassic to Cretaceous age in uplifted Paleozoic-Mesozoic terranes. Significant Cenozoic placers accumulatedwhere Paleozoic gold lodes that formed along the active mar-gin of Gondwana have been uplifted and eroded, such as Vic-toria, Australia, Westland, New Zealand, and the easternCordillera, South America. Similar giant placers apparentlydid not form where lodes of the same age, but hosted in Pre-cambrian basement rocks, were uplifted and eroded, as in theNorth China craton.

There is a gap in gold placers between 2.0 Ga and 60 Ma(Fig. 5). The significant Paleoproterozoic placer deposits ofTarkwa (Ghana) and Jacobina (Brazil) accumulated 100 to200 m.y. prior to orogenic gold lodes of the same terranes, sig-nifying earlier gold-forming episodes (Groves et al., 2005).Orogenic gold provinces may also have been eroded from thecatchment to the Witwatersrand basin (Frimmel et al., 2005),although opinion remains divided as to the origin of thosegold deposits (Law and Phillips, 2005). Preservation of Pre-cambrian placer gold deposits is a function of the thick,buoyant, and refractory continental lithospheric mantle thatpreserved ancient foreland basins, as well as the Mesozoic-Cenozoic Cordilleran-style accretionary orogens in whichorogenic gold provinces developed.

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Metallogeny of Continent-Continent Orogens

Metallogenic context

Continent-continent orogens, such as the Tethyan Alpine-Himalayan, Damaran, or Appalachian-Caledonian, are un-likely to be highly metalliferous. Given little production of ju-venile crust, many of the Au-, Hg-, and Sb-rich veins thatdevelop in accretionary orogens do not form during rework-ing of older continental terranes that already have been highlydevolatilized. Also, given the single narrow suture zone insuch orogens, there is a narrow magmatic arc and thus a lim-ited extent to the distribution of arc-related epithermal, por-phyry, and skarn deposits. Continent-continent orogens maybe the end stages of accretionary orogens, where an externalocean closes (Sengor and Natal’in, 1996a), such as where theHimalayan orogen has followed the Altaid orogen.

Therefore, mineral deposits that typify older accretionaryorogens are commonly preserved in the trapped juvenilecrust during subsequent collisional orogenesis (e.g., the cen-tral Asian deposits between the Tethysides blocks and Siber-ian craton). Podiform Cr and VMS deposits that formed inoceanic arcs later become emplaced as ophiolite fragmentswithin obducted sheets during continent-continent collision.

Granitoid Sn-W

Magmatic deposits of Sn-W are associated with granites incontinent-continent–type orogens, as well as some Cordilleranand Andean orogens. Source magmas are the highly fraction-ated peraluminous granites of Ishihara’s (1981) ilmenite series,involving melting of reduced sedimentary facies and mantlemelts. These granites are enriched in the incompatible ele-ments Cs, Rb, Th, U, Nb, Ta, Sn, W, Mo, and LREE, and thevolatile elements F and B (Sinclair, 1996). Intrusions and min-eralization are constrained by structures imposed by regionaltectonics (Clarke et al., 2000). Mineralization is triggered bymixing of saline magmatic and low-salinity meteoric watersduring regional uplift (Kontak and Clark, 2002).

Large Sn-W metallogenic provinces developed in Paleozoic-Mesozoic continent-continent orogens. Prominent depositsare Grey River and East Kemptville, Nova Scotia, in the Ap-palachian orogen; Xhuashan, China; the Erzgebirge and Mas-sif Central, provinces at 325 to 300 and 290 to 260 Ma, and thesouthwest England and Portugal (Panasqueira) province at~290 Ma of the Variscan orogen in Europe; and the Sinobur-malaya terrane of the Permo-Triassic (300–200 Ma) orogen ofsoutheastern Asia (Pollard et al., 1995). Granites were gener-ated in overthickened crust following collision and emplacedin a tensional regime, possibly after delamination of continen-tal lithospheric mantle and gravitational collapse of crust.

Prominent granitoid Sn-W provinces of inner arcs of An-dean orogens are Llallagua, Chojlla, and Chambillaya in theTertiary of Bolivia, and San Rafael and Pasto Bueno of Peru.Cordilleran-type orogenic Sn-W provinces include the Aber-foyle and Ardlethan districts of the Tasman orogen, as well asthe Regal Silver and Kalzas deposits in northwestern NorthAmerica (Sinclair, 1996).

The conjunction of elements that lead to the formation oforogenic Sn-W granites are (1) siliciclastic sediments from aweathered catchment deposited under the chemocline, (2)input of mantle melts, (3) melting of overthickened crust, and

(4) advection of fluid-undersaturated melts to shallow crustallevels. According to Clark et al. (1990), the richest granite-re-lated Sn-W deposits in the Andes reflect an arc-broadeningevent, caused by a shift in subduction angle linked to chang-ing convergence velocity. The secular distribution of Sn-Wprovinces is similar to that of porphyry Cu deposits, includingpreservation potential in eroding mountain belts (Fig. 5).

Metallogeny of Foreland Basins

U: Foreland-intracontinental basins

Thirty percent of the global U resource is sited in Protero-zoic siliciclastic sequences, proximal to unconformities (Fig.5; Ruzicka, 1996). Sedimentary basins evolved on all cratonsafter supercontinent assembly at ~2.0 to 1.8 Ga (Windley,1995), but preserved economic deposits of U have only beenfound in foreland-intracratonic basins (Fig. 4B, D) in NorthAmerica, Australia, and western Africa. Nash et al. (1981)suggest that the U geochemical cycle was widely establishedfrom 2 Ga onward, with Proterozoic sedimentary accumula-tions setting the stage for 1.0 Ga metamorphic U provinces inthe Grenville and Damara orogens, as well as the sedimen-tary-hosted U provinces of western Texas and the ColoradoPlateau.

The conjunction of two geodynamic events may have beenresponsible for initiating large-scale near-surface U geo-chemical cycles; i.e., transition from flat to steep subductionat ~2.6 Ga, and decreased intensity of plume activity since 2.6Ga (except the 1.9 Ga superplume; Figs. 3A, 5). Archean slabmelt TTG possess Th/U ratios of 5.8 and lower U abundancesthan younger intermediate to felsic arc magmas, where Th/Uratios average 3.7 (Drummond et al., 1996). Lower plumeintensity increased the continental freeboard, permitting de-velopment of extensive continental siliciclastic sequences(Fig. 3B).

The Athabasca basal unconformity developed on Archeancratons and in Paleoproterozoic metasedimentary rocks thatinclude reductants such as Fe2+ and graphite. The unconfor-mity corresponds to a transition from intracontinentaltranstensional basins controlled by escape tectonics outboardof the Trans-Hudson orogen, to a Laramide-type distal fore-land or perimeter basin (cf. Dickinson et al., 1988), to theTrans-Hudson orogen. Collision of the Reindeer and Hearneterranes from 1810 to 1710 Ma generated topographic upliftof ~10 km in the Trans-Hudson orogen. Four depositional se-quences total 1,800 m of quartzose sandstones capped by 500m of oolitic or stromatolitic dolomite, with three regional un-conformities; these sequences developed in a fluvial to lacus-trine environment with an aeolian input (Ramaekers andCatuneanu, 2004). Sedimentation commenced at 1830 Ma inthe Athabasca and correlative Thelon basin (Rainbird et al.,2002), continuing intermittently to 950 Ma. Coeval cratonicsequences in Laurentia include the Hornby Bay Group, andSioux, Baraboo, and Mazatzal sandstones (Ross, 2000).

Two meteoric water-dominated hydraulic systems devel-oped during extensional subsidence at ~1500 Ma. Isotopicallyevolved fluids at 180º to 240ºC advected through the base-ment, interacting with reductants, whereas formation brinesdissolved U as aqueous U+6 within the Athabasca sequence.Uranium was deposited where the two fluids mixed proximal

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to the unconformity, where it was transected by northeast-trending faults that parallel the structural grain of Trans-Hud-son orogen accreted terranes (Kotzer and Kyser, 1993, 1995).This event was coeval with assembly of the Colombia super-continent. A second mineralization stage developed at ~1400Ma, coeval with intracontinental rifting of Columbia; a thirdstage at 1260 Ma was related to extension during the Macken-zie large igneous province event; and a fourth occurredduring diking in the Athabasca basin or possibly in responseto distal tectonic events, such as Nipigon rifting, the Racklanorogen in the Yukon, or the Grenville orogen (Kotzer et al.,1992; Kyser et al., 2000; Ramaekers et al., 2005).

In Australia, the McArthur River foreland basin developedduring the 2.0 to 1.8 Ga Barramundi orogen. As much as 15km of siliciclastic and carbonate sediments accumulated from1800 to 1770 Ma in a marine to terrestrial environment withintermittent volcanism. The principal unconformity-relateddeposits are Jabiluka, Ranger, and Nabarlek, in the PineCreek sub-basin of the McArthur River basin. Uraninite pre-cipitated at 1640 Ma from saline (Na-Mg-Ca-Cl), diageneticfluids >100 m.y. after termination of sedimentation, as in theAthabasca basin. A pronounced change in the apparent pale-omagnetic wandering path throughout the McArthur Riverbasin at 1640 Ma corresponds to the timing of both sedimen-tary rock Pb-Zn-Ag and U mineralization (Idnurm et al.,1995). Diagenetic fluids advected through the basin intermit-tently for >900 m.y. (Kyser et al., 2000, and referencestherein; Polito et al., 2004).

There are several common factors in the evolution of theAthabasca and McArthur basins and their U deposits, as wellas the 2 Ga Oklo U province, Gabon. These include (1) Pale-oproterozoic orogens that sutured supercontinents, such asBaltica, Laurentia, and East Antarctica into Nena (Fig. 8A;Rogers, 1996; Rogers and Santosh, 2004); (2) reductants inthe basement; (3) foreland basins that promoted high hy-draulic conductivity in sediments, (4) evaporites in some se-quences that generated saline diagenetic brines; and (5) evo-lution to intracontinental basins underlain by some amount ofArchean continental lithospheric mantle, which accounts fortheir preservation compared to geodynamically equivalentPhanerozoic basins. Protracted fluid flow was tectonicallytriggered and generated multiple stages of mineralization>100 m.y. after sedimentation (Hoeve and Quirt, 1987; Ra-maekers et al., 2005).

Rollfront sandstone-hosted deposits on most continentsrepresent ~30 percent of global U resources; they formed at<100 Ma (Fig. 5B). Key factors in their formation are (1) de-velopment of extensive upland terrestrial forests at ~100 Ma;(2) intermontane or intracratonic basins with fluvio-lacustrinesediments characterized by the conjunction of large hydraulicconductivities, with both oxidized and reduced facies; (3) tec-tonic uplift induced orographic rainfall; and (4) topographi-cally driven fluid flow (Nash et al., 1981).

Carbonate-hosted Pb-Zn

Carbonate-hosted MVT deposits of Pb-Zn are reviewedby Leach et al. (2005a). These workers emphasize that oro-genic uplift is a key factor to create elevated recharge fortopographically driven flow of formation fluids throughsedimentary rock aquifers in foreland basin sequences (but

see Cathles and Adams, 2005, for an alternative explanationin terms of gas-driven flow). Unconformities developed at theforebulge (Fig. 4D) tend to be favorable sites for discharge ofmetalliferous fluids or mixing of fluid reservoirs.

The MVT deposits are restricted to compressional tectonicevents of Paleozoic and Cretaceous-Tertiary age. The eco-nomically most significant ores precipitated in the Devonianto Permian assembly of Pangea, notably Appalachian orogen-related deposits of the United States mid-continent. YoungerMVT deposits formed during the North American Cordilleranorogen (Robb Lake, Canada) and the Africa-Eurasian Tethyanorogen (Cracow-Silesian, Poland). Accumulation and preser-vation of platform sedimentary sequences that include car-bonates are fundamental factors in the spatial and secular dis-tribution of these huge base metal deposits.

Metallogeny of Intracontinental Rifts

Fe oxide-Cu-Au-Ag-REE

The Fe oxide-Cu-Au deposit class has a variety of metalbudgets, possibly reflecting a spectrum of crustal depths(Hitzman et al., 1992; Davidson and Large, 1998; McMaster,1998; Porter, 2002; Williams et al., 2005). Economically themost important deposits in decreasing age include Carajas,Brazil (2.57 Ga); Kiruna-Bergsdalen, Sweden (1890–1880Ma); the Great Bear magmatic zone, Canada (1885–1865Ma); the Cloncurry District of the Mt. Isa terrane (1.79–1.74Ga) and Olympic Dam (1.59 Ga), Australia; and the St.Francois district of Missouri (1450–1350 Ma). Variants mayinclude Slipfontein, an Fe-Cu-Au-F deposit pipelike deposit,hosted by a 2.06 Ga Bushveld granite; the 2.09 Ga Palaboracarbonatite ring complex, sited at the margin of the Kaapvaalcraton (Figs. 5, 8B); and Kiruna as an Fe-dominant end mem-ber (Pirajno, 2000; Groves and Vielreicher, 2001). The unify-ing characteristics are enrichment of Fe-P-F and alkali meta-somatism of the host rocks (Pirajno, 2000; Groves et al.,2005).

The Olympic Dam Cu-Au-Ag-REE deposit developedwithin the 1.59 Ga Roxby Downs A-type granite. Lithos-pheric attenuation was focused at the eastern margin of theArchean Gawler craton where the Torrens hinge zone devel-oped at the transition to thinner post-Archean continentallithospheric mantle. The Adelaide intracontinental rift basinfilled with a siliciclastic and volcaniclastic rock sequence,most likely with evaporites. Basaltic magmas were generatedby decompressional melting of hot asthenosphere and/or amantle plume that advected into the base of thinned lithos-phere. Basalts ponded at the Moho, fusing refractory, halo-gen-rich, lower crust, the residue of previous hydrous meltextraction, to form anhydrous A-type granites emplaced atshallow crustal levels (Campbell et al., 1998). Two fluidsmixed in the breccia: a cooler, oxidized, hypersaline formationbrine at hydrostatic pressure carrying Cu-Au-U-S, and a litho-statically pressured deeper Fe-F-Ba-CO2-rich fluid likelyevolved from A-type magmas (Haynes et al., 1995).

Deposits of the St. Francois Mountains are associated withA-type felsic to intermediate anorgenic magmas. The εNdvalues of LREE-enriched mineralization in these deposits ofProterozoic to Tertiary age span depleted mantle to enrichedvalues in common with associated igneous rocks, indicating

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the mixing of basaltic with fused lower crustal reservoirs anda magmatic source of REE (Gleason et al., 1999).

The interplay of geodynamic and geologic elements forthis deposit class are (1) attenuated continental lithosphere,(2) basaltic liquids at the base of the crust and a rift sequencenear the surface, (3) translithospheric rift-related faults tofocus A-type magmas, (4) enhanced hydraulic conductivity atshallow crustal levels by hydraulic fracturing, and (5) mixingof two fluids with different redox states. The 1.8 to 1.1 Gaage span for most of these deposits is related to the assemblyand dispersal of the supercontinent Columbia and then as-sembly of Rodinia (Fig. 5; Unrug, 1997). Their location nearattenuated continental lithospheric mantle may involve alka-

line magmatism generated by decompressional melting ofmetasomatized continental lithospheric mantle. Many Pro-terozoic sedimentary basins feature evaporites. Groves et al.(2005) list possible Phanerozoic counterparts, such as Cande-laria, with the caveat that these may have characteristics in-termediate between the Proterozoic Cu-Au-REE-Fe depositsand Au-rich porphyry systems.

Sedimentary rock-hosted Pb-Zn

The dominantly Proterozoic SEDEX mineralization oc-curred in intracontinental rifts that developed into passivemargins (Fig. 4), as reviewed by Leach et al. (2005b). Barleyand Groves (1992) related the secular distribution of these

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FIG. 8. A. Configuration of the Mesoproterozoic supercontinent Columbia of Zhao et al. (2004), illustrating the distribu-tion of Paleo- to Mesoproterozoic SEDEX Pb-Zn deposits proximal to Archean or Paleoproterozoic margins. Recast fromLydon (2000). B. Columbia, after Zhao et al. (2004), with 1.8 to 1.5 and 1.4 to 1.1 Ga belts of anorogenic magmatism fromHaapla and Rämö (1999).

deposits to the fragmentation stage of supercontinents. In thereconstruction by Zhao et al. (2004) of the Mesoproterozoicsupercontinent Columbia, the giant Pb-Zn deposits ofAustralia, Laurentia, and India are on intracontinental riftsthat presaged drift, proximal to the margins of Archean cra-tons or Paleoproterozoic terranes (Fig. 8A). The regionallyextensive, and protracted, history of diagenetic brine flow ofthese Proterozoic siliciclastic basins, reflected both in uncon-formity U and SEDEX Pb-Zn deposits, results from the com-bination of basin development on relatively thick continentallithospheric mantle with the thermal anomaly of the 1.4 Gasuperplume that dispersed Columbia. Sulfides precipitatedfrom basinal brines pumped by far-field forces, which wereactive from synsedimentary times to ~200 m.y. after sedi-mentation, e.g., Sullivan and HYC, respectively (Idnurm etal., 1995). Extensive passive margin sequences did not formprior to the Paleoproterozoic given low freeboard (Fig. 3).According to Large et al. (2001), younger intracontinentalrifts lack such deposits.

Sedimentary rock-hosted Cu-Co deposits

Copper sulfide and native Cu deposits, commonly withhigh Co concentrations and locally PGE (Kucha, 1982; Hitz-man et al., 2005), occur as stratiform deposits in fine-grainedclastic sedimentary rocks and silty dolomites within rift-re-lated sedimentary basins. The best known of these sediment-hosted Cu deposits occur in the Central African Copperbeltof the Democratic Republic of the Congo and Zambia (Selleyet al., 2005) and the Kupferschiefer of central Europe. Iso-lated large deposits include White Pine (Michigan), Redstone(Northwest Territories, Canada), and Dzhezkazgan (Kaza-khstan). These deposits are of various, commonly impreciselyconstrained, ages, ranging from Neoproterozoic (CentralAfrican Copperbelt, Redstone, White Pine), through Car-boniferous (Dzhezkazgan), to Permian (Kupferschiefer) butshare a common tectonic setting in failed intracratonic rifts oraulacogens (Fig. 5; Raybould, 1978).

Many of these rifting events have been terminated by basininversion, resulting in deformation, metamorphism, and latehydrothermal overprinting (e.g., Richards et al., 1988a, b).These effects have hampered interpretations of original met-allogenic processes, to the extent that models ranging fromsynsedimentary deposition (Renfro, 1974; Garlick, 1981),through early diagenetic fluid flow (Bartholomé et al., 1973),to late epigenetic hydrothermal Cu introduction (Sales, 1962)have all been proposed over the last several decades (see re-views by Gustafson and Williams, 1981; Kirkham, 1989;Sweeney et al., 1991; Selley et al., 2005).

In the case of the Central African Copperbelt, the timing ofsedimentary and tectonic events is not well constrained.However, it is clear that synsedimentary or diagenetic modelswould require a pre-Lufilian orogeny age (pre-600–550 Ma;Porada and Berhorst, 2000), whereas an epigenetic modelwould suggest an early or syn-Lufilian age, because the oresare clearly deformed by this orogenic event. Richards et al.(1988a) obtained a two-stage Pb-Pb model age for least re-crystallized Cu-Fe sulfides from the Musoshi deposit of 645 ±15 Ma, which was interpreted either to reflect the timing ofCu introduction into the Ore Shale, or isotopic disturbance ofpreexisting sulfides. Either explanation indicates that primary

mineralization predated the main stage of Lufilian orogene-sis. Selley et al. (2005) similarly report an unpublished Re-Osdate of 816 ± 62 Ma for stratiform sulfide deposition at theKonkola deposit in Zambia, consistent with a diagenetic orlate-diagenetic timing for mineralization. Richards et al.(1988a, b) also dated rutile and uraninite from late quartzveins cutting the Ore Shale at Musoshi and obtained an ageof 514 Ma, indicating that these veins postdated the Lufilianorogeny, and that they could, therefore, not have been re-sponsible for original introduction of Cu into the Ore Shale.

Ore textures indicate a permeability control on metal dis-tribution, reflected in the concentration of Cu sulfides withinmore sandy laminae of the siltstone sequence, with finergrained, less permeable laminae being almost devoid of sul-fides except possibly syndepositional pyrite (Richards et al.,1988b). Permeability in the sediments during introduction ofCu implies an origin prior to regional metamorphism, per-haps during early diagenesis, because pore space would sub-sequently have been filled by diagenetic and then metamor-phic minerals (Brown, 1978). An early diagenetic timingwould also coincide with advanced development of the riftbasin, involving crustal thinning and increased mantle-de-rived heat flow. In contrast, Selley et al. (2005) propose amodel involving secondary permeability development duringearly orogenic fluid flow.

A common characteristic of sedimentary rock-hosted Cudeposits is that they are hosted by what were originally or-ganic-rich black shales or dolomites, typically representingthe first marine transgression in previously subaerial clasticsedimentary basins (Oszczepalski, 1999). In some Cu-Coprovinces, such as the Central African Copperbelt, the un-derlying clastic rock sequences are referred to as red beds, re-flecting subaerial oxidation of relatively immature sandstones(commonly dune-bedded), arkoses, and conglomerates; evap-oritic horizons also occur locally (Jackson et al., 2003). Knife-sharp contacts between the Ore Shale and underlying con-glomerates attest to sudden flooding of the subaerial basin,with a switch to deposition of fine-grained clastic sedimentaryrocks with high organic content, followed by deeper marinecarbonate deposits. A link is thus suggested between hydro-carbon maturation, diagenetic flow of oxidized basinal brinesin the footwall sequences, and base metal sulfide depositionby reduction upon interaction between these brines and or-ganic-rich shales (Annels, 1979; Kelly and Nishioka, 1985;Sverjensky, 1987; Jowett, 1992; Mauk and Hieshima, 1992).Red-bed formation has been suggested as a key precursor fac-tor in this process, by causing the breakdown of primary sili-cate and oxide minerals to render trace concentrations of baseand other metals labile (Zielinski et al., 1983; Brown, 1984).These metals are then available for dissolution by later fluxesof warm, oxidized basinal brines (Rose, 1976).

Expulsion of metalliferous brines from the deeper parts ofsedimentary basins is recognized to be an essential part of theore-forming process, not only of sedimentary rock-hosted Cudeposits but also of other sedimentary rock-hosted base metaldeposits such as MVT and SEDEX Pb-Zn deposits (Cathlesand Adams, 2005; Leach et al., 2005a). The different metalinventory of these deposits, but otherwise similar environ-ments of formation within intracratonic sedimentary basins,may reflect simply the dominant composition of sedimentary

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rocks. For example, Sverjensky (1989) has suggested that Pb-and Zn-rich deposits form where brines have flowed predom-inantly through sandstone and carbonate aquifers, respec-tively, whereas Cu-rich deposits form where aquifers containa significant amount of immature sediments, such as red-bedarkoses. The high Co contents of some Central African Cudeposits may reflect leaching from mafic materials, eitherpresent as clastic components in the arkosic sediments or sillsdeep within the sedimentary sequence (Annels and Sim-monds, 1984).

The onset of basin inversion is commonly regarded as an im-portant tectonic driving force for fluid flow, resulting in highfluid pressures in deeper parts of the basin that force fluids toescape by percolation through normally impermeable shalehorizons, thereby bringing oxidized metalliferous brines into di-rect contact with reductants in these shales (Cathles and Adams,2005). Metal deposition as sulfides also requires a source of re-duced sulfur, which may be generated by in situ reduction ofsulfate carried by the same brines (McGowan et al., 2003).

The formation of sedimentary rock-hosted Cu deposits, likeother sedimentary basin-hosted base metal deposits, may thusbe seen as part of the larger supercontinent cycle, formingduring the early stages of rifting (Raybould, 1978; Barley andGroves, 1992; Titley, 1993). Successful rifting will generate anew ocean basin, with the original rift sediments forming partof a passive margin sequence. However, such sequences areeither currently submarine, or have been caught up in andpotentially destroyed by later collisional events, and aretherefore either inaccessible to, or of low potential for, explo-ration. In contrast, failed rifts have high preservation poten-tial within stable continental interiors and are thus the mostprospective regions for discovery of economic deposits.

Passive Margins

Phosphorites

Most sedimentary phosphate deposits accumulated on thecontinental shelves of the western margins of continents andin passive margin marine settings, within 45° of paleoequa-tors. Deposition occurred in zones of high bioproductivityfrom upwelling of cold polar currents moving toward theequator in oceanic gyres. Ocean basins ~3,000 km wide arerequired for gyres, implying deposition of phosphate 15 to 20m.y. after rifting. Deposits may also form on east-facing pas-sive margins, such as in the Miocene basin of Florida (Fig.4C; Chandler and Christie, 1996).

Significant phosphorite units were deposited as the first ex-tensive passive margins developed during dispersal of the su-percontinent Kenorland at ~2.4 to 2.2 Ga (Fig. 5). Examplesare units in the 1.95 to 1.85 Ga Animikie Group in Minnesotaand in the 2.0 Ga Trans-Amazonian Central Guiana belt, Suri-name. Major phosphate accumulations became widespreadon passive margins (e.g., Russian platform) following breakupof the supercontinent Rodinia at ~600 to 500 Ma, whichmarked the Neoproterozoic-Cambrian boundary. The largestdeposit is the Permian (300–251 Ma) Posphoria Formation,Montana and Idaho, deposited on the western margin of latePaleozoic Pangea in an epicratonic sea. These deposits are as-sociated with global sea-level high stands linked to maxima ofplume or ocean-ridge activity. Phosphates also occur on ocean

plateaus and islands (Windley, 1995; Chandler and Christie,1996; Follmi, 1996; Ilyin, 1998).

Placer Ti-Zr-Hf

Economically the most significant placer deposits of tita-nium minerals (Garnett and Bassett, 2005) and zircon, in ter-ranes associated with gemstones, are those of eastern andWestern Australia, South Africa, southwestern India, and thesoutheastern United States. An economic example is the <5Ma beach sands of the Swan Coastal Plain, Western Australia.This plain is bounded to the east by the Darling fault scarp,linked to the dispersal of Pangea at ~230 Ma. Zircon age pop-ulations in the beach sands and Cretaceous siliciclastic rocksof the Perth basin are both 1.3 to 1.1 Ga and 600 to 500 Ma.The catchment is thought to be Pan-African terranes, such asthose of the Albany-Fraser or Pinjarra orogens, involved inwelding the Australian, Antarctic, and Indian continents intoPangea. Those orogens are interpreted to have been formerlylocated to the west, possibly partially in India prior to sub-duction under Asia. Titanium minerals and zircon were con-centrated during recycling of the Cretaceous sedimentaryrocks (Freeman and Donaldson, 2004, and referencestherein). Tertiary equivalents are present on the western pas-sive margin of South Africa, which are sourced in the Meso-proterozoic Namaqua province (Macdonald et al., 1997).

Sedimentary rock-hosted Pb-Zn

Neoproterozoic-Phanerozoic sedimentary rock-hosted Pb-Zn deposits, in contrast to older Proterozoic examples, are re-stricted to Atlantic-type passive margins (Leach et al., 2005a,b). An essential factor in evolving ore-forming brines is thepresence of sabka sediments within carbonate platforms, ashas been suggested for the Red Dog deposit, Alaska. MostSEDEX deposits formed within 30° of the paleoequator(Leach et al., 2004, 2005a, b).

Metallogeny of Anorogenic Magmatic Belts

Rapakivi Sn

More than 70 percent of Proterozoic anorogenic magma-tism occurs in a 5,000-km-long belt, which is 1,000 km wide,extending from southern California through Labrador to theSvecofennides. Windley (1995) recognized six stages in theassembly and dispersal of the Mesoproterozoic superconti-nent Columbia and subsequent formation of the Grenvilleorogen: (1) accretion of arc terranes of the Yavapai, CentralPlains, Penokean, and Svecofennide provinces from 1.9 to 1.8Ga; (2) minimum melts of hydrated and accreted terranes togenerate 1.85 to 1.81 Ga granites; (3) melting of overthick-ened lower crust that yielded vapor-poor nonminimum melts,such as the 1.8 Ga Svecofennian granites, and lamprophyresfrom enriched domains of mantle wedge; (4) delamination ofoverthickened continental lithospheric mantle replaced byhot asthenosphere, which generated 1.75 to 1.55 Ga anoro-genic Rapakivi granites from lower crust and gabbros from as-thenosphere melts; (5) plume activity that thinned and frag-mented the supercontinent at ~1.3 Ga, with basaltic liquidsponding at the Moho density filter, fractionating to anorthosite-gabbro complexes; and (6) the Grenville ocean opened andclosed in an Alpine-Himalayan-type orogen, leaving remnants

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of anorogenic Rapakivi granites and anorthosites (Fig. 8B;Windley, 1995, figs. 15.2, 15.3; Karlstrom et al., 2001).

Rapakivi granites are anorogenic A-type granites, a mem-ber of bimodal gabbro-anorthosite and granitoid complexes(Frost et al., 2002). Most Rapakivi provinces formed from 1.6to 1.0 Ga, with a tectonic setting of a mantle plume imping-ing on thinned, incipiently rifted continental lithosphere orupwelling of asthenosphere following extensional collapse, asin the Basin and Range province. Rapakivi granites are meltsof residual lower crust and are relatively anhydrous and re-duced; accordingly, they crystallized at shallow crustal levels.Tin deposits with Be, W, Zn, and Cu are present in late-stagedifferentiates of 1.7 to 1.5 Ga Rapakivi granite complexes ofFennoscandia, Missouri, and Brazil (Haapala and Rämö,1999; Pirajno, 2005).

Anorthosite Fe-Ti-V

Dominantly Mesoproterozoic in age, these deposits arewithin layered or massive anorthosite-gabbro complexes (Fig.5). Plume-related basalts ponded at the Moho density filterwhere there was extensive fractional crystallization of plagio-clase. Associated magmas are jotunites and monzonorites thatare rich in Fe, Ti, V, and P; apatite suppresses crystallizationof magnetite, resulting in immiscible Fe-rich evolved liquids.Orebodies are cumulus layers or discordant bodies of Fe andTi oxides, the latter crystallizing from immiscible liquids.

Prominent examples are Kiglapait, Canada; Smaalands-Taberg, Sweden; and Kachkanar-Kusinskoye, Russia (Gross andScoates, 1996). In the southern hemisphere, a 300-km-long beltbetween Angola and Namibia includes the 2.0 Ga Kunene In-trusive Complex with an Fe-Ti-V province (Pirajno et al., 2004).

The narrow secular duration for massif anorthosites andtheir associated Fe-Ti-V accumulations is readily accountedfor by plume-continental lithospheric mantle interactions.Plumes impinging on thick Archean continental lithosphericmantle did not undergo decompressional melting. After the1.9 Ga superplume event, mantle plume activity waned butwas sufficiently hot during dispersal of the Mesoproterozoicsupercontinent to evolve large quantities of basalts under ex-tended post-Archean continental lithospheric mantle. Shear-wave splitting is greatest at the margins of Archean continen-tal lithospheric mantle, and Vp and Vp/Vs ratios of theGrenville are consistent with 20-km-thick, plagioclase-richlower crust (Musachio and Mooney, 2002; Sleep et al., 2002).

Plume-Continental Lithosphere Interaction

Magmatic Ni-Cu-PGE and stratiform chromite

Three types of Ni-Cu sulfide and PGE deposits are recog-nized (Arndt et al., 2005; Barnes and Lightfoot, 2005;Cawthorne et al., 2005): (1) Ni-Cu sulfide accumulations aspart of ultramafic-mafic intrusive complexes and continentalflood basalts (CFB) within intracontinental rifts, with Ni/Curatios <1; (2) komatiite-hosted deposits with Ni/Cu ratiosgenerally >10; and (3) tholeiitic intrusions in greenstoneterranes or along translithospheric faults characterized byNi/Cu ratios of 2 to 3 (Naldrett, 1989; Eckstrand, 1996). Allthree reflect plume-lithosphere interaction (Fig. 2B).

The Stillwater Complex, Montana, with chromite and Ni-Cu sulfide domains, is an intrusive expression of the 2.7 Ga

global superplume event (Fig. 3A; Isley and Abbott, 1999; Pi-rajno, 2000). The Great Dyke, emplaced at 2050 Ma, repre-sents extension in the Zimbabwe craton, following amalgama-tion of the Kaapvaal and Zimbabwe cratons. The 2060 to 2050Ma Bushveld Intrusive Complex was emplaced proximal tothe Murchison-Thabuzimbi lineament that also controlledthe architecture of the intracratonic Transvaal sedimentarybasin. These African cratons were possibly adjacent to theAntarctic and Pilbara cratons at ~2.5 to 2.2 Ga (Pirajno,2000). The intrusive complexes are likely an early stage of the1.9 Ga superplume (Fig. 3A). In the Great Dyke andBushveld Intrusive Complex, oxide ores of Cr, Ti, Fe, and Vand sulfide ores of Ni, Cu, Co, and PGE were associated withultramafic liquids possessing high Mg number but low in-compatible element abundances (Pirajno, 2000, 2005). Giventhe constraint on depth of plume decompressional meltingimposed by the thick continental lithospheric mantle,Archean intrusive complexes such as the Bushveld may eitherreflect lateral flow of plume melts into the craton or, alterna-tively, hotter Archean plumes melted at greater depths (Xie etal., 1993).

Critical factors for transition metal ores in ultramafic tomafic magmatic bodies are an increase of SiO2 content to in-duce S saturation and open-system conditions. Increase ofSiO2 content of the parental liquid occurs either by assimila-tion of crustal rocks or by mixing with noritic melts (Fig. 2B).In an open system, sulfides equilibrate with successive pulsesof melts or by mixing of melts. Many Archean and Proterozoicmafic-ultramafic intrusive complexes have vast quantities ofnorites. Norites are not evolved, or crustally contaminated,tholeiitic basalts. Intriguingly, these intracontinental noritesfeature incompatible element enrichment in conjunctionwith depletions of Nb-Ta, the characteristics of convergentmargin mafic magmas (Hall and Hughes, 1990; Pearce andPeate, 1995). Some intrusive complexes have units with U-shaped REE patterns compositionally akin to Phanerozoicboninites and recent boninites of the Izu-Bonin-Mariana arc(Stern et al., 1991; Taylor et al., 1994). Shallower mantlelithosphere of Archean terranes acquired a subduction zonesignature in subcreted normal oceanic and ocean plateaulithosphere during accretionary assembly of the terranes intocratons. Subsequently, the deep residue of plume meltingcoupled buoyantly to form the deeper continental lithos-pheric mantle (Wyman et al., 2002; Schmitz et al., 2004),which later remelted at shallower depths by decompressionduring extension and/or plume impingement. This generatedintracratonic norites with a subduction signature and allowedmixing of plume material with high Si norite liquids in layeredcomplexes (Fig. 2B).

Other well-documented examples of magmatic Ni-, Cu-,and Co-bearing sulfide deposits stemming from plume im-pingement on incipiently rifted lithosphere are discussedbelow. In the circum-Superior craton belt, Ni sulfide depositsin Manitoba occur at the cratonic margin, in ~1.8 Ga sillscompositionally evolved from dunite to pyroxenite, which arean expression of the ~1.9 Ga superplume event (Fig. 3A;Condie et al., 2001). The 1850 Ma Sudbury igneous complexis located at the boundary between the Archean Superior cra-ton and Proterozoic Southern province. Large volumes ofnorite are present, hosting Cu-, Ni-, and Co-bearing sulfides

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with significant PGE. Crustal melting was induced by a me-teorite inpact (Barnes and Lightfoot, 2005), so this igneouscomplex is an exception to the association of magmatic Ni-Cusulfide ores with mantle plumes. In the Neoproterozoic, Ni-Cu sulfide deposits occur in (1) the 1.1 Ga Duluth complex ofthe mid-continent rift associated with the Keweenawan largeigneous province; (2) the 1.1 Ga Coppermine large igneousprovince of the Northwest Territories, with the Muskox in-trusive complex; and (3) the Jinchuan deposit, China, whichoccurs in ultramafic bodies that intruded translithosphericfaults at the southwestern margin of the North China craton.Deposits of the Insizwa complex, South Africa, are associatedwith the Karoo large igneous province and formed at 200 to180 Ma; impingement of the ancestral Iceland plume on Lau-rentia-Baltica induced their rifting, as part of the TertiaryNorth Atlantic large igneous province. In Greenland, the 55Ma Skaergaard intrusion, hosting PGE-Au deposits, is a relictof that plume-lithosphere interaction (Fig. 3A; Saunders etal., 1997; Pirajno, 2000).

Magmatic Ni-Cu-Co-PGE deposits in the Noril’sk-Talnakhmetallogenic province have clearcut expressions of the cou-pled geodynamic and magmatic elements that are associatedwith this deposit type. The province is sited at the edge of theSiberian craton, where the transition from thick (Archean) tothinner continental lithospheric mantle guided the location ofthe regional, translithospheric Kharayelakh fault. Incipientrifting created intracontinental basins between the Siberian,eastern European, and Taimyr cratons. Impingement of aplume at 250 Ma near the failed triple junction led to exten-sive decompressional melting under thin continental lithos-pheric mantle, and plume magmas erupted onto a Devonianepicontinental sedimentary sequence generating 3.5-km-thick continental flood basalts. Tholeiitic basalts are preva-lent, with minor alkali basalts and picrites indicating meltingin an anomalously hot plume tail. Assimilation of low S conti-nental crust led to increase of SiO2 content and S saturationof basaltic melts, with gravitational accumulation of magmaticsulfides that partitioned Ni-Cu-PGE from multiple pulsesthrough open-system magma conduits. More than 12 Gt of Sentered the system from stoping of sulfate-rich evaporites,but only ~1 percent of this S entered the orebody (Naldrett,1989; Lightfoot and Hawkesworth, 1997).

The largest komatiite-hosted Ni-Cu deposits are in the 2.7Ga Norseman-Wiluna belt, Yilgarn craton. Komatiite flowserupted in a deep marine environment over sulfidic sedi-ments deposited in a ~200-km-wide intracontinental rift. Sul-fur saturation of the ultramafic liquids may stem from assim-ilation of the sediments (Lesher and Keays, 2002). Similardeposits are present in the 2.7 Ga Abitibi belt andNeoarchean greenstone terranes of Botswana and Zimbabwe.Paleoproterozoic equivalents formed in greenstone terranesof Manitoba, Ontario, Quebec (Ungava, Raglan), Finland(Hitura), Russia (Pechenga), and Tanzania (Kabanga). A com-mon geodynamic element for these deposits is eruption ofhigh-temperature, S-undersaturated ultramafic melts throughcontinental (Noresman-Wiluna) or dominantly oceanic(Abitibi) crust (Naldrett, 1989; Eckstrand, 1996; Cassidy etal., 2002; Lesher and Keays, 2002).

Tholeitic intrusion-hosted cumulus Ni-Cu sulfide depositsoccur dominantly in Archean greenstone terranes, with fewer

in Paleoproterozoic counterparts of Finland and Russia (Nal-drett, 1989; Ekstrand, 1996). Their common secular distribu-tion with komatiite-hosted deposits is consistent with plumemagmas advecting into an arc, back-arc, or sub-continentalflood basalt setting.

Magmatic Ni-Cu-Co sulfide deposits at Voisey’s Bay, New-foundland, are sited on the 1.8 Ga translithospheric suturebetween the Archean Nain and Paleoproterozoic Churchillprovinces. Troctolite magmas, likely part of the 1350 to 1290Ma anorogenic Nain Plutonic Complex, intruded the sutureat 1.3 Ga, coeval with dispersal of the supercontinent Colum-bia, triggered by a mantle plume. Interaction with graphite-bearing paragneisses of the host terrane by assimilation andfractional crystallization added Si, K, Na, and S to the trocto-lites, triggering S saturation and segregation of immisciblesulfide liquids (Eckstrand, 1996; Naldrett and Ripley, 2001).In summary, magmatic Ni deposits have the same secular dis-tribution as mantle plumes.

Economic stratiform chromite deposits are all Archean orPaleoproterozoic in age. The largest deposits are Selukwe, inthe 3420 Ma Sebakwian sequence of the Zimbabwe craton;Kemi in Finland (2444 Ma); and Campo Formoso in Brazil(2000 Ma). All involve plumes interacting with Archean con-tinental lithospheric mantle (Fig. 2B; Duke, 1996b). Strati-form chromite in the Neoarchean Bird River Sill, Manitoba,and Big Trout Lake intrusion, Ontario, appear to be the resultof plume-related intrusions emplaced into Archean green-stone terranes (Duke, 1966b). Crystallization of chromite wastriggered by mixing of a high Mg primitive melt with SiO2-rich norites, raising the Si activity in the former.

Diamonds

Diamonds form by reaction of asthenospheric carbonatiticliquids with peridotite (p-type) and eclogite (e-type) of deep,mostly Archean, continental lithospheric mantle (Gurney etal., 2005). Accordingly, ages of inclusions in diamonds span3.3 Ga to Mesoproterozoic (Fig. 2B; Kirkley et al., 1991). Car-bon is introduced into the continental lithospheric mantleboth from deep asthenospheric fluids and from subductedocean crust, in keeping with independent evidence for resi-dence of subducted material in Archean continental lithos-pheric mantle (Cartigny, 2005). Diamonds are transported asxenocrysts from the continental lithospheric mantle to shal-low crustal levels in kimberlites or lamproites, both incom-patible element-enriched and volatile-rich ultramafic mag-mas (Mitchell, 1995; Dawson, 1999). Kimberlitic melts aregenerated in the upper mantle, some at depths of 450 to 670km as indicated by inclusions of beta majorite garnet, but mayalso form below the 670-km D' transition zone. In southernAfrica, continental lithospheric mantle with slower P-wavevelocity correlates with a greater proportion of eclogitic sili-cate inclusions in diamonds, younger Sm-Nd ages of the in-clusions, more depleted δ13C, and fewer diamonds character-ized by low N contents. Converse properties characterizehigh P-wave domains of continental lithospheric mantle(Shirey et al., 2004). Whereas mantle plumes do not undergodecompressional melting at ~300 km beneath Archean conti-nental lithospheric mantle, volatile-rich kimberlites possessthe buoyancy flux to penetrate this mantle along preexistingstructures (Fig. 2B).

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Three of seven known major kimberlite events are associ-ated with superplumes: (1) ~480 Ma in Russia, China,Canada, South Africa, and Zimbabwe; (2) ~280 Ma in Lau-rentia-Baltica; and (3) ~120 to 80 Ma in North America,India, Siberia, Brazil, and Africa, linked to the Pacific Creta-ceous superplume and associated dispersal of Gondwana(Figs. 3A, 5). Jelsma et al. (2004) identified four lineamenttrends in southern Africa, along which many kimberlitesoccur, that they attributed to lithospheric structures formedduring breakup of Gondwana. Oceanic lithosphere, stored at670-km depth, may avalanche to the core-mantle boundary,ejecting superplumes that in turn cause dispersal of super-continents (Condie, 2002). Kimberlite events not known tobe associated with plumes occurred at 1 Ga, 410 to 370 Ma,200 Ma, and 50 Ma (Condie, 2001). Geodynamic settings ofkimberlites are reviewed by Helmstaedt (1993); plume-lithosphere interaction is prevalent, but continental rifts andtransform faults are also significant for localizing kimberliteemplacement in the crust.

Iron formations

Arguably the most significant insight into the fundamentalprocess for iron formations comes from the work of Isley andAbbott (1999), reviewed by Clout and Simonson (2005). Theydemonstrated that from 3.8 to 1.9 Ga, iron formations andocean plateaus that were erupted from mantle plumes have acommon time series (Fig. 9; see also Clout and Simonson,2005). Reduced hydrothermal fluids enriched in Fe2+ and Si,from convection through submarine basaltic lavas, weretransported by ocean circulation to shallower basins where Feprecipitated in near-surface waters. Fryer et al. (1979) pio-neered the concept of large volcanic-related hydrothermalfluxes into Archean oceans, specifically to maintain anArchean CO2 greenhouse. Simonson and Hassler (1996) ar-gued for deposition of Archean banded iron formation (BIF)below the wave base in deep water during global sea-levelhigh stands, in keeping with decreased continental freeboardassociated with oceanic plateaus (Fig. 3).

Isley and Abbott’s (1999) insight explains the scarcity ofiron formations younger than 1.8 Ga (Fig. 9). Accordingly,these deposits not only span the putative great oxygenationevent at ~2.2 Ga and, therefore, are not proxies for the oxida-tion state of Earth’s atmosphere-hydrosphere system but re-quire oxygenated waters to precipitate Fe3+ (Ohmoto,2004a,b). Corroborative evidence for this depositionalscheme of reduced source fluids and oxygenated surface ma-rine waters comes from 2.9 Ga BIF in India, which are char-acterized by positive Eu but negative Ce anomalies. The for-mer is indicated by the solubility of Eu2+ in reducedhydrothermal fluids, whereas the latter is consistent with se-questration of Ce3+ from marine water by Fe3+ and Mn4+ ox-ides and oxyhydroxides (Kato et al., 2002). Not all Precam-brian iron formations have such clear Eu and Ce anomalies.

Trendall and Blockley (2004) reject the conventional classi-fication of iron formations into Algoman and Superior types.They identify four main associations. The first is older BIF involcanic basins of Archean greenstone terranes of the Slave,Superior, Baltic, Ukraine (Krivoi Rog BIF), Dharwar, Ama-zon, Yilgarn, Kaapvaal, and West African cratons. Microbandsare interpreted as chemical and seasonal varves, and BIF are

associated with graywackes. Similar BIF occur in a 650 Maarc terrane of the Pan African orogen and in the OrdovicianBathurst district (Windley, 1995; Peter, 2003). From the geo-chemistry of first-cycle volcanogenic turbidites in the 2.7 GaAbitibi greenstone terrane, BIF were precipitated in posi-tions distal from bimodal intraoceanic arcs or in back arcs, aswell as distal from ocean plateaus (Feng et al., 1993). On fourof the Gondwana continents there is a peak of areally exten-sive BIF at 2.5 to 2.4 Ga. These BIF unconformably overlythe Archean Karnataka, Amazon, Transvaal, and Pilbara cra-tons. The Hammersly and Transvaal BIF may have been con-tiguous in a superbasin termed Vaalbara (Fig. 9; Cheney,1996). Associated rocks are mafic volcanics and black shales.An equivalent, the 2.3 Ga Kursk district, is present on themargin of the Siberian craton (Trendal and Blockley, 2004).Molecular fossils in Hammersly BIF confer evidence for pho-tosynthetic cyanobacteria, and concentrations of P, V, Co, Zn,and Mo in these BIF are consistent with precipitation by Fe2+

oxidizing bacteria, as also characterizes present-day Fe-richaqueous environments (Brocks et al., 1999; Kornhauser et al.,2002). Shales of the Transvaal Supergroup are cratonic(Wronkiewicz and Condie, 1990), and the areal extent of Pa-leoproterozoic BIF on the Gondwana continent is consistentwith stable continental shelves above Archean continentallithospheric mantle.

Granular iron formations (GIF) accumulated on circum-Superior province continental shelves in the Lake Superiorand Labrador regions, and in the Earaheedy basin on thenorthern margin of the Yilgarn craton, at ~1.9 to 1.8 Ga (Pi-rajno, 2005). These GIF are commonly associated with epi-clastic sedimentary rocks and tuffs; they are interpreted as re-working of Fe oxide particles in a shallow-water, high-energyenvironment (Simonson, 2003; Trendall and Blockly, 2004).The areal extent of GIF stems from stable continental shelveson, or proximal to, Archean continental lithospheric mantle.

Distinctive iron formations, termed Rapitan, were de-posited from ~800 to 700 Ma. The principal accumulationsare Rapitan in the Yukon, Urucum in Brazil, and in theDamara belt, Namibia (Fig. 9C). These iron formations areassociated with, but more restricted than, Neoproterozoicglacial deposits and include dropstones. Evidence for rift-re-lated mafic magmatism that generated Fe-rich hydrothermalplumes is present in the stratigraphic sequences (Yeo, 1986;Young, 1988; Trompette et al., 1998). A possible analogue isFe-rich hydrothermal sediments in the Red Sea, where rift-ing is caused by the African superswell and related plumes.Iron formations are also associated with the ~250 Ma andCretaceous superplumes (Fig. 9; Oyarzun et al., 2003).

The combination of factors necessary for iron formationsare (1) mantle plumes; (2) tectonic stability for timescales of>1 m.y.; (3) hiatus of proximal volcanic activity over compara-ble timescales; (4) basin architecture that promoted open ex-change with deep marine bottom waters; and (5) sufficientwater depth to limit the input of epiclastic sediments, GIF ex-cepted (Isley and Abbott, 1999; Trendall and Blockly, 2004).

SynthesisPlume intensity was relatively greater in the Archean and

erupted hotter melts, but some type of plate tectonics wasalso operating. Archean cratons formed where ocean plateaus

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that had been erupted from plumes became jammed againstconvergent margins and the buoyant refractory residue of hotplume melting coupled with imbricated arc-plateau crust.This refractory residue constitutes the deep continentallithosphere mantle keel that defines Archean cratons and isresponsible for preservation of Archean continental crust andits deposits. Magmatic Ni deposits are associated with ko-matiites and basalts erupted from mantle plumes; VMS de-posits formed in intraoceanic arcs; and orogenic gold depositsare prevalent in the Neoarchean (Fig. 10), linked toCordilleran-type orogens that welded cratons into the first su-percontinent, Kenorland, at ~2.7 Ga.

The Proterozoic is characterized by a distinctive set of min-eral deposits. At deeper crustal levels, magmatic Ni-Cu de-posits formed in layered complexes where high Mg melts from

mantle plumes intruded translithospheric structures guided bythe transition from thicker Archean to thinner Proterozoic con-tinental lithospheric mantle. At shallower depths, Fe oxide Cu-Au-REE deposits are also controlled by structures marginal toArchean continental lithospheric mantle. As plume intensitywaned, the continental freeboard increased, and phosphorites,carbonates, and Fe and Mn formations precipitated on the firstextensive passive margins as Kenorland dispersed. The first Uaccumulations were in foreland basins to orogens that weldedColumbia; the first Pb-Zn deposits in intracontinental rifts ac-companied the dispersal of Columbia; and anorthosite-associ-ated Fe-Ti-V and Rapakivi Sn deposits occur in the vast belts ofProterozoic anorogenic magmatism that fundamentally reflectplume-lithosphere interactions. Several features are evident inFigure 10: The sparsity of deposits that form in topographically

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FIG. 9. (A). Secular distribution of mantle plumes, after Isley and Abbott (1999). (B). Iron formations after Trendall andBlockley (2004). Rapitan (C), Algoman (D), and VMS (E) after Ohmoto (2004a). (F). Volume of ocean crust from Condie(1997).

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FIG. 10. Secular variation of specified classes of mineral deposits according to geodynamic setting. Peak height on the y-axis is scaled according to relative size of the metallogenic provinces. A. AM = anorogenic magmatism; CA = continental arc;CC = continent-continent orogen; CO = Cordilleran orogen; CR = continental rift; IA = intraoceanic arc; PL = plume-lithos-phere. Porphyry-epithermal and VMS deposits form in both intraoceanic and continental arcs, but for simplicity of illustra-tion the former are plotted on the continental arc track. Similarly, magmatic Sn deposits occur in both Cordilleran andcontinent-contenent orogens, but are illustrated only on the latter. B. Sedimentary basins. BA = back arc; FA = fore arc; FL= foreland; IC = intracontinental; O = oceanic; PM = passive margin; RM = rifted continental margin; SS = strike slip. Placergold deposits accumulate in the fore arcs and back arcs of orogenic belts, but for simplicity of illustration are plotted in forearcs. Sources: Meyer (1988), Goldfarb et al. (2001) for orogenic Au, Groves et al. (2005) for Fe oxide-Cu-Au-REE.

elevated tectonic belts, such as magmatic Sn and porphyry-ep-ithermal deposits, in the Archean; the sparsity of several deposittypes over the interval from ~1.8 to ~0.8 Ga; the onset of sev-eral classes of sedimentary rock-hosted deposits with the firststable passive margins and increased freeboard; prevalence ofFe-Ti-V deposits in belts of Proterozoic anorogenic magmatism;and the low prospectivity of intracontinental settings.

In terms of preservation, the sparsity of many deposit typesfrom ~900 to 500 Ma may have resulted from a secular de-crease in thickness and buoyancy of the continental lithos-pheric mantle, coupled with Grenvillian orogens having deeplevels of erosion due to delamination of continental lithos-pheric mantle. The secular distribution of ore deposits in thePhanerozoic (Fig. 10) reflects enhanced preservation, espe-cially of deposits in topographically elevated ranges, notwith-standing thinner continental lithospheric mantle.

Four potential future directions for research may provideuseful insights for exploration. At the scale of cratons, betterseismic imaging of continental lithospheric mantle topographymay assist in the exploration for magmatic Ni-Cu and Fe oxideCu-Au-REE deposits. Refined reconstructions of the super-continent cycle allow projections of metallogenic provinces(Fig. 8). At the scale of terranes, investigations on the conjunc-tion of thermal, structural, and lithological factors will help todetermine the distinction between a metallogenic provinceversus regions of subdued mineralization. At the scale of aprovince, efforts to systematize Damkohler (NdD) numbers(Johnson and DePaolo, 1994) will help to determine why largeor small deposits of a given type may form from the same ore-forming fluids but with subtleties of geochemistry that may in-dicate size; e.g., large deposits may have high Nd signatures.

AcknowledgmentsWe are grateful to Bruce Eglington, Franco Pirajno, Paul

Ramaekers, Vlad Sopuk and Derek Wyman for reviewingsome, or all, sections of an intial draft of this manuscript. Thesection on geodynamics draws on a document written by AliPolat and RK for an unpublished report to the Canadian As-sociation of Mining Industry Research Organization(CAMIRO). Economic Geology One Hundredth AnniversaryVolume reviewers, Dallas Abbott and David Groves, conferredinsights and identified errors that resulted in substantial im-provement to the final version. Glen Caldwell, Kevin Cassidy,Bruce Eglington, and Mike Lesher guided RK to informationwhere background was lacking. Karen McMullan and IgnacioGonzales are thanked for assistance with the text, and RyanSchmidt, June McLintock, and Tim Wardell for generating thefigures. RK acknowledges the George McLeod endowment tothe Department of Geological Sciences at the University ofSasktchewan, and JPR and RK acknowledge support of Dis-covery Grants from the Natural Sciences and Engineering Re-search Council of Canada. We appreciate the invitation by JeffHedenquist to write this article.

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