Ore Geology Reviews 70 (2015) 323–345
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Tectonic, magmatic, and metallogenic evolution of the Tethyan orogen: From subduction to collision Jeremy P. Richards ⁎ Dept. Earth and Atmospheric Sciences, University of Alberta, Edmonton, Alberta T6G 2E3, Canada
a r t i c l e
i n f o
Article history: Received 1 April 2014 Received in revised form 14 November 2014 Accepted 20 November 2014 Available online 27 November 2014 Keywords: Paleotethys Neotethys Metallogeny Tectonics Magmatism Subduction Collision Ore deposits
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http://dx.doi.org/10.1016/j.oregeorev.2014.11.009 0169-1368/© 2014 Elsevier B.V. All rights reserved.
a b s t r a c t This paper reviews the tectonic, magmatic, and metallogenic history of the Tethyan orogen from the Carpathians to Indochina. Focus is placed on the formation of porphyry Cu ± Mo ± Au deposits, as being the most characteristic mineral deposit type formed during both subduction and collisional processes in this region. Relatively little is known about the history of the Paleotethys ocean, which opened and closed between Gondwana and Eurasia in the Paleozoic, and few ore deposits are preserved from this period. The Neotethyan ocean opened in the Permian–Early Triassic as the Cimmerian continental fragments (the cores of Turkey, Iran, Tibet, and Indochina) rifted from the northern Gondwana margin and drifted northwards. These microcontinents docked with the Eurasian margin at various points in the Mesozoic and Cenozoic, and formed a complex archipelago involving several small back-arc basins and remnants of the Paleotethyan ocean. The main Neotethyan ocean and these smaller basins were largely eliminated by collision with India and Africa–Arabia in the early Eocene and earlymid Miocene, respectively, although Neotethyan subduction continues beneath the Hellenic arc and the Makran. The majority of porphyry-type deposits are found in association with Neotethyan subduction (mainly in the Mesozoic and Paleogene), and syn- to post-collisional events in the mid-Paleogene to Neogene. They are found throughout the orogen, but some sections are particularly well-endowed, including the Carpathians–Balkans– Rhodopes, eastern Turkey–Lesser Caucasus–NW Iran, SE Iran–SW Pakistan, southern Tibet, and SE Tibet– Indochina. Other sections that appear barren may reflect deeper levels of erosion, young sedimentary cover, or lack of exploration, although there may also be real reasons for low prospectivity in some areas, such as minimal subduction (e.g., the western Mediterranean region) or lithospheric underthrusting (as proposed in western Tibet). Over the last decade, improved geochronological constraints on the timing of ore formation and key tectonic events have revealed that many porphyry deposits that were previously assumed to be subduction-related are in fact broadly collision-related, some forming in back-arc settings in advance of collision, some during collision, and others during post-collisional processes such as orogenic collapse and/or delamination of subcontinental mantle lithosphere. While the formation of subduction-related porphyries is quite well understood, collisional metallogeny is more complex, and may involve a number of different processes or sources. These include melting of: orogenically thickened crust; previously subduction-modified lithosphere (including metasomatized mantle, underplated mafic rocks, or lower crustal arc plutons and cumulates); or upwelling asthenosphere (e.g., in response to delamination, slab breakoff, back-arc extension, or orogenic collapse). The most fertile sources for syn- and post-collisional porphyry deposits appear to be subduction-modified lithosphere, because these hydrated lithologies melt at relatively low temperatures during later tectonomagmatic events, and retain the oxidized and relatively metalliferous character of the original arc magmatism. Unusually metallically enriched lithospheric sources do not seem to be required, but the amount of residual sulfide phases in these rocks may control metal ratios (e.g., Cu:Au) in subsequent magmatic hydrothermal ore deposits. Relatively Au-rich deposits potentially form in these settings, as observed in the Carpathians (e.g., Roşia Montană), Turkey (Kisladag, Çöpler), and Iran (Sari Gunay, Dalli), although the majority of syn- and post-collisional porphyries are Cu–Mo-rich, and resemble normal subduction-related deposits (e.g., in the Gangdese belt of southern Tibet). This similarity extends to the associated igneous rocks, which, being derived from subduction-modified sources, largely retain the geochemical and isotopic character of those original arc magmas. While still retaining a broadly calc-alkaline character, these rocks may extend to mildly alkaline (shoshonitic) compositions, and may display adakite-like trace element signatures (high Sr/Y and La/Yb ratios) reflecting melting of deep crustal garnet amphibolitic sources. But they are otherwise hard to distinguish from normal subduction-related magmas.
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Small, post-collisional mafic, alkaline volcanic centers are common throughout the orogen, but for the most part appear to be barren. However, similar rocks in other post-subduction settings around the world are associated with important alkalic-type porphyry and epithermal Au ± Cu deposits, and the potential for discovery of such deposits in the Tethyan orogen should not be overlooked. © 2014 Elsevier B.V. All rights reserved.
1. Introduction The Paleotethys and Neotethys ocean basins formed in the Paleozoic and Mesozoic, respectively, between the Laurentian/Laurasian continental masses to the north, and the Gondwana continents to the south. Plate tectonic reconstructions for the Paleotethys are the subject of considerable debate, and models from, for example, C.R. Scotese (PALEOMAP Project, www.scotese.com) and Stampfli and Borel (2004) disagree in many details such as subduction polarity and the locations of individual continental fragments. Nevertheless, there is general agreement that the Paleotethys ocean basin was first formed in the mid-Paleozoic, and was then progressively destroyed by convergence between Laurentia and Gondwana, culminating in formation of the supercontinent Pangea in the late Paleozoic (Fig. 1A). Stampfli and Borel (2004) indicate that by the Late Triassic, only small remnants of the Paleotethys remained. Meanwhile, the Neotethyan ocean basin had begun forming in the Permian–Early Triassic by rifting of the Cimmerian continental fragments from the northern margin of Gondwana (Fig. 1A). These fragments include the cores of present-day Turkey, Iran, Tibet, and Indochina, which swept northwards as the Paleotethys closed, eventually to accrete to the southern margin of Laurasia in the Late Triassic–Early Jurassic. From this point forward, the history of the Neotethys involves northward subduction below the accretionary Laurasian margin (Figs. 1B and 2), the opening of small back-arc basins along that margin (e.g., the Pindos and Vardar oceans), and eventual (ongoing) closure by collision with Africa–Arabia and India (Fig. 2D). Relatively few known mineral deposits are convincingly associated with the Paleotethys ocean, and its geological record is not well preserved. In contrast, the record is much better for the Neotethys, and numerous world class mineral deposits are associated with its formation and closure. Consequently, the focus of this paper is largely on the Neotethys ocean, and in particular Neotethyan subduction- and collision-related magmatism on the Laurasian margin. Porphyry Cu ± Mo ± Au and related epithermal Au ± Cu deposits are the predominant mineral deposit type in the orogen, and, while some of these deposits appear to be related to normal subduction-related magmatism (e.g., Clark and Ullrich, 2004; von Quadt et al., 2005), there has been increasing recognition that many are related to post-subduction collisional processes (e.g., C.R. Harris et al., 2013; Harangi et al., 2007; Heinrich and Neubauer, 2002; Hou et al., 2003, 2004; Hou et al., 2009; J.X. Li et al., 2011; Janković, 1997; Lu et al., 2013a,b; Moritz et al., 2010; Neubauer, 2002; R. Wang et al., 2014a, 2014b, 2014c; Richards, 2009; Richards et al., 2006; Roşu et al., 2004; Shafiei et al., 2009; X.-S. Wang et al., 2014; Yang et al., 2009). The wide range of tectonic settings represented along the Neotethyan orogenic belt thus provides a good opportunity to study porphyry ore formation in response to different geodynamic processes. The Neotethyan orogen stretches for over 12,000 km from the Alps, through the Carpathians–Balkans, Turkey, Iran, and Pakistan, Tibet, and Indochina, and includes sections where Neotethyan oceanic lithosphere is still being subducted (the eastern Mediterranean and the Makran) to advanced continental collision (the Alps and Himalayas). In order to organize the presentation of material in this review, I have separated the orogen into three main sections, based on their predominant geodynamic context: (1) the collided arcs of the Carpathians and Balkans; (2) the incipient Afro-Arabian collision zone of Turkey–Iran– Western Pakistan; and (3) the advanced Indian collision zone of the
Pamir–Himalayas–Indochina. The Alpine section of the orogen is excluded because few significant porphyry deposits are located in this part of the belt, likely due to deeper levels of erosion and smaller total volumes of subduction in the western Tethys. The text is kept intentionally brief, and tectonic reconstructions are simplified where possible, so that focus can be maintained on the high-level processes controlling magmatism and ore formation at different times and places along the belt. An extensive bibliography is provided where more detailed information can be found. 2. Sources of data In order to visualize the spatio-temporal information being reviewed here, I have compiled paleogeographic reconstructions of the Tethyan region from the Cretaceous to Neogene, using maps generated from the Ocean Drilling Stratigraphic Network's Plate Tectonic Reconstruction Service (www.odsn.de/odsn/services/paleomap/ paleomap.html; Fig. 2). I also reproduce, with permission, two paleogeographic maps for the Permian and Late Jurassic from C.R. Scotese's PALEOMAP Project (Scotese, 2007; Fig. 1). For present-day locations of major Tethyan suture zones and ore deposits, I have synthesized a map from numerous sources that are indicated in the caption to Fig. 3. In several cases, especially for the traces of older sutures, there is disagreement within the literature. Consequently, in attempting to introduce some consistency and continuity along the length of this extensive belt, it has been necessary to take some liberties with individual published interpretations. The maps shown should therefore not be considered definitive for any given region, but merely illustrative of the general structure of the overall belt. Figs. 2 and 3 also show the locations of selected porphyry-type deposits, drawn mainly from the databases of Singer et al. (2005, 2008), and color coded by broad age group. Major deposits (with N~ 100 Mt resource) are named in Fig. 3, with their ages shown where known; smaller deposits in Singer's databases are shown as smaller symbols, and are mostly not named for clarity. Readers are referred to these databases for grade and tonnage figures, which are not repeated or updated here because these data change rapidly and are readily available for most deposits on the internet. 3. Carpathians and Balkans 3.1. Mesozoic In the Mesozoic, the Carpathians and Balkans were located close to a hinge zone between the Laurasian and Gondwanan continental masses, such that the Neotethyan ocean was never very wide at this point (Figs. 1B and 2A). As the central Atlantic ocean began to open in the Jurassic (Fig. 1B), the Gondwanan continental block rotated anticlockwise northwards towards the eastern Laurasian margin, and the Carpathian–Balkan region was characterized by microcontinent and arc collisions, and the opening and closure of small, likely backarc basins such as the Vardar ocean (Fig. 2). The history of these basins is widely debated (e.g., Channell and Kozur, 1997): Robertson et al. (2013b) suggest that the Vardar ocean formed in the Late Triassic– Early Jurassic between the Korabi-Pelagonian and Serbo-Macedonian continental blocks, whereas Stampfli and Borel (2004) indicate that this ocean basin was partially a remnant of the Paleotethys, and was expanded by back-arc extension in the Jurassic. Subduction of the
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Fig. 1. A. Paleogeography of the Paleotethys ocean in the Late Permian; note the incipient rifting of the Cimmerian continental fragments from the northern Gondwana margin to form the Neotethys ocean. B. Paleogeography of the Neotethys ocean at its maximum extent in the Late Jurassic; initiation of rifting to form the central Atlantic ocean at this time, followed by south Atlantic rifting in the Cretaceous, resulted in anticlockwise rotation and northward drift of Africa–Arabia (and later India) to progressively close the Neotethys ocean in a scissor-like motion around an axis close to the western Mediterranean. Images reproduced with permission from Scotese (2007). Abbreviation: CIMM = Cimmerian continental fragments (parts of Turkey, Iran, Tibet, Indochina).
Vardar ocean to the northeast beneath the Serbo-Macedonian continent in the Late Jurassic led to arc magmatism in the Rhodopes dated from 164 to 155 Ma (Fig. 2A; Anders et al., 2005; Jahn-Awe et al., 2010). The Vardar ocean finally closed in the Late Cretaceous–early Cenozoic, resulting in collisional tectonics (Fig. 2B, C; Robertson et al., 2013b). A number of large porphyry Cu–Au and related high-sulfidation Au deposits were formed in association with Late Cretaceous calc-alkaline
arc magmatism in this region (Ciobanu et al., 2002; Clark and Ullrich, 2004; Janković, 1997; Lips, 2002), which has been variously termed the Bananitic magmatic and metallogenic belt (BMMB; Ciobanu et al., 2002) or the Apuseni–Banat–Timok–Srednogorie belt (ABTS; von Quadt et al., 2005). Major porphyry Cu–Au deposits include: Moldova Nouă in Romania; Majdanpek, Veliki Krivelj, and Bor in Serbia; and Elatsite and Assarel in Bulgaria (Fig. 3). The large Chelopech highsulfidation epithermal Au deposit in Bulgaria is spatially associated
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Although most researchers agree that these porphyry systems and their generative calc-alkaline magmas were related to subduction of the Vardar ocean (e.g., Neubauer, 2002), there is disagreement in detail, reflecting the uncertainties of plate tectonic reconstructions in this structurally complex region. Von Quadt et al. (2005) and Zimmerman et al. (2008) describe the Late Cretaceous magmatism in terms of trench
with porphyry magmatism (Chambefort et al., 2007), and the nearby Elatsite porphyry is unusual in having elevated platinum group element (PGE) concentrations (Augé et al., 2005; Eliopoulos et al., 2014; Tarkian et al., 2003). These deposits mostly have ages ranging from 92 to 84 Ma, except for Moldova Nouă (65 Ma) (ages are from Singer et al., 2005, 2008, and references therein).
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Fig. 2. Paleogeographic reconstructions of the Neotethyan region at (A) 145.5, (B) 100 Ma, (C) 50 Ma, and (D) 15 Ma. These reconstructions are redrawn from maps generated using the Ocean Drilling Stratigraphic Network's Plate Tectonic Reconstruction Service (www.odsn.de/odsn/services/paleomap/paleomap.html). Plate motions are modeled relative to a magnetic reference frame, using a Mercator projection. Blue lines represent present-day coastlines, for reference. Locations of major porphyry deposits forming around the time of each image are approximate; see Fig. 3 for more precise locations, deposit names, and ages. Abbreviations: A = Afghan block; C = Carpathians; CI = Central Iranian block; K = Kirşehir block; L = Lut block; M = Moesian Platform; P = Pontides; R = Rhodopes; SA = South Armenian block; SSZ = Sanandaj–Sirjan Zone; TAB = Tauride–Anatolide block.
J.P. Richards / Ore Geology Reviews 70 (2015) 323–345 Fig. 3. Topographic relief map of the Alpine–Himalayan Tethyan orogenic belt, showing major structures, Tethyan sutures, and porphyry Cu ± Au deposits grouped by age; larger deposits (generally N100 Mt resource) are identified by name, with age where known, and larger symbol sizes. Locations and ages of porphyry deposits are derived principally from Singer et al. (2005, 2008), with updated information from Richards et al. (2006, 2012), Perelló et al. (2008), Taghipour et al. (2008), and Imer et al. (2013). Tethyan sutures and structures are derived primarily from Stampfli and Kozur (2006), with additional information from Wortel and Spakman (2000), Badarch et al. (2002), Metcalfe (2006, 2013), Piper et al. (2006), Yin (2006), Robinson et al. (2007), Zhang et al. (2010), Yakubchuk et al. (2012), Pirajno (2013), and Deng et al. (2014). Plate velocities relative to Eurasia from Calais and Amarjargal (2000), Guillot et al. (2003), Allen et al. (2004), Vernant et al. (2004), and Regard et al. (2005). The background topographic relief map was generated by Hans Braxmeier as a layer for Google maps, and is available from Maps-For-Free.com. Abbreviations: EAF = East Anatolian Fault; IAES = İzmir–Ankara–Erzincan suture zone; NAF = North Anatolian Fault; SSZ = Sanandaj–Sirjan Zone; UDMA = Urumieh–Dokhtar magmatic arc.
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retreat (slab rollback) and orogenic collapse during oblique subduction, whereas Chambefort and Moritz (2006) suggest that the Panagyurishte ore district in Bulgaria formed in a transtensional strike-slip fault system, possibly in a back-arc setting or in response to slab breakoff. The plate reconstructions shown in Fig. 3A and B certainly imply highly oblique convergence along the Serbo-Macedonian margin at this time, which may have contributed to the formation of favorable transtensional or transpressional structural loci for shallow crustal magma emplacement and porphyry formation (e.g., Tosdal and Richards, 2001). 3.2. Paleogene Closure of the Vardar ocean in the Paleogene resulted in a shift of subduction to the Hellenic trench, the onset of extension in the Aegean, and a switch to collisional tectonics along the former Balkan arc (Dumurdzanov et al., 2005; Georgiev et al., 2010; Jolivet and Brun, 2010; Jolivet et al., 2013; Kaiser-Rohrmeier et al., 2013). Postsubduction or collision-related magmatism with minor associated Late Eocene–Oligocene porphyry and epithermal mineralization occurred in the Balkans (e.g., Buchim; Fig. 3) and eastern Rhodopes (Heinrich and Neubauer, 2002; Janković, 1997; Moritz et al., 2010). Marchev et al. (2005) relate these Paleogene ore deposits in the Rhodopes to late orogenic extension and metamorphic core complex formation, perhaps triggered by lithospheric mantle delamination (Schefer et al., 2011). The Recsk porphyry Cu, skarn, and epithermal deposits in the Carpathians of Hungary also formed during this period (35.5 Ma; Baksa et al., 1980; Baksa, 1986; Singer et al., 2005, 2008; K–Ar ages of ~ 28 Ma are reported for adularia and illite from low sulfidation epithermal parts of the system by Molnár et al., 2008). Molnár (2007) relates the generative calc-alkaline magmatism at Recsk to oblique collisional processes as the Adriatic microplate contacted the European margin.
4. Afro-Arabian collision zone The Afro-Arabian collision zone runs through Turkey, Armenia/ Azerbaijan, Iran, and western Pakistan. Although the geology and tectonic history are broadly similar along this length of the Tethyan orogen, they are sufficiently different in detail to merit subdividing the following descriptions by country, while attempting where possible to correlate across international borders. The level of documentation also varies by country, with the most detailed information being available for Turkey, followed by Iran, and least for Pakistan and Armenia/Azerbaijan. Key references are provided below to support detailed geological and tectonic events, but the following sources are also used as general references: central Tethyan paleogeography: Şengör and Yilmaz (1981), Dixon and Robertson (1984), Dercourt et al. (1986), Şengör (1987), Stampfli (2000), Stampfli and Borel (2004), Moix et al. (2008), and Robertson et al. (2013a); Turkish geology and mineral deposits: Bozkurt and Mittwede (2001), Boztuğ et al. (2003), and Yigit (2006, 2009); Iranian geology and mineral deposits: Berberian and King (1981), and Richards (2003b). 4.1. Paleotethys The geological record of the Paleotethys is slight in this region. Cambrian calc-alkaline granites occur in the Bitlis Massif of southeastern Turkey, and are thought to have formed in response to Paleotethyan subduction beneath the northern margin of Gondwana, prior to separation of the Cimmerian continental fragments (Ustaömer et al., 2009). In contrast, Early Carboniferous granites from the eastern Pontides of Turkey do not have subduction-related geochemical signatures, although they are generally thought to be related to northward subduction of the Paleotethys beneath the southern Eurasian continental margin (Ustaömer and Robertson, 2010). No significant porphyry deposits are reported in association with these Paleozoic rocks.
3.3. Neogene
4.2. Mesozoic
The distinctive curvature of the Carpathians is thought to have developed in response to post-collisional deformation and rotation since 13 Ma (Ciobanu et al., 2002; Csontos et al., 1992; Dupont-Nivet et al., 2005; Nemcok et al., 1998). Seghedi et al. (1998, 2004) and Neubauer et al. (2005) suggest that the key tectonic process was rollback of the remnant intra-Carpathian oceanic slab as the Adriatic microplate collided with the European foreland. Several important ore deposits formed in the Carpathians at this time, including the Roşia Poieni porphyry Cu–(Au) deposit, and low-sulfidation epithermal Au deposits such as Roşia Montană and Sacarimb (Heinrich and Neubauer, 2002; Manske et al., 2006). Miocene calc-alkaline magmatism in the Carpathians is not thought to be directly related to subduction, but was generated during postcollisional extension by partial melting of subduction-modified lithospheric mantle and crust (Harangi et al., 2007; Neubauer et al., 2005; Roşu et al., 2004). C.R. Harris et al. (2013) specifically suggest that lithospheric metasomatism occurred during Mesozoic Neotethyan subduction, and that this subduction-modified material was remobilized by collisional processes to generate fertile magmas. Of significance also is the Miocene formation of the Skouries Pt–Pd– Au-rich porphyry Cu deposit on the Chalkidiki peninsula of northern Greece (19 Ma; Frei, 1995). This deposit is unusual in terms of its high Au and PGE contents (Eliopoulos and Economou-Eliopoulos, 1991; Eliopoulos et al., 2014), and its shoshonitic affinity (Kroll et al., 2002). The geodynamic setting for this deposit is not clear but it appears again to be post-subduction or collision-related (Economou-Eliopoulos and Eliopoulos, 2000) and as such it may fit the model of precious metal-enriched post-subduction porphyry deposits proposed by Richards (2009).
The central part of the Tethyan belt now occupied by Turkey, Iran, and western Pakistan is made up of a collage of island arcs and continental fragments (Cimmerian continents) originally rifted from the northern margin of Gondwana in the Permian (Fig. 1). These fragments (including Tibet and parts of Indochina) collided with the southern Eurasian margin in the Late Triassic–Early Jurassic along the Paleotethys suture (Figs. 1 and 2). Remnants of the Paleotethys ocean may be preserved in the Caspian and Black Sea basins (e.g., Eyuboglu et al., 2011, 2012), although localized back-arc rifting in the Mesozoic and Cenozoic confuses this picture. 4.2.1. Turkey Closure of the Paleotethys ocean in the Early Jurassic sutured the Pontides to the Eurasian margin, and was followed by initiation of northward subduction of a branch of the Neotethys beneath the newly accreted margin (Fig. 1B; Dokuz et al., 2010). Early Jurassic to Late Cretaceous calc-alkaline granitoid intrusions in the Pontides are related to this period of subduction (Boztuğ et al., 2006; Kaygusuz et al., 2008; Ustaömer and Robertson, 2010), but mostly appear to be eroded below the levels of preservation of porphyry and epithermal deposits. Numerous ophiolites were accreted to the southern Pontide margin during this period, and a small back-arc basin, the Artvin Basin, opened in the eastern Pontides in the Early–Middle Jurassic (Ustaömer and Robertson, 2010). Back-arc spreading may also have occurred in the western Black Sea basin in the Late Cretaceous (Espurt et al., 2014; Okay et al., 2013) or earlier (Zonenshain and Le Pichon, 1986). Collision of the small Kırşehir micro-continental block (also called the Central Anatolian Crystalline Complex) with the Central Pontide margin along the İzmir–Ankara–Erzincan suture zone (Fig. 3) occurred
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in the Late Cretaceous to Paleocene (Espurt et al., 2014; Kaymakci et al., 2009; Lefebvre et al., 2013; Meijers et al., 2010; Robertson et al., 2009). However, the paleogeographic reconstructions shown in Fig. 2 suggest that the Kırşehir block may have only been separated from the Pontides by a small back-arc basin, and collision may have been relatively “soft”. The more substantial (or “hard”) collision of the Tauride–Anatolide block (TAB; Fig. 3) in the late Paleocene–early Eocene is discussed in Section 4.3. Late Cretaceous–Eocene calc-alkaline magmatism on the southern margin of the TAB (SE Anatolia) is thought to reflect northward subduction of the main Neotethyan ocean along the Bitlis–Zagros subduction zone (Parlak, 2006; Rızaoğlu et al., 2009; Robertson et al., 2007; Şengör and Yilmaz, 1981; Yilmaz, 1993). 4.2.2. Lesser Caucasus Relatively little is known about the geology of the Lesser Caucasus region of Armenia and Azerbaijan, where deformation was caused by collision between the South Armenia block with Eurasia in either the Late Cretaceous (Rolland et al., 2009) or Paleocene (Sosson et al., 2010). The paleogeographic reconstruction shown in Fig. 2 suggests that a back-arc basin linked to the Black Sea opened between these blocks in the Cretaceous, and progressively closed in the Cenozoic. Several porphyry Cu–Mo deposits of Jurassic, Cretaceous, and Paleogene age occur in the South Armenian block, including Tekhut (Figs. 2 and 3), while other deposits such as Agarak formed in the Eocene. Mederer et al. (2014) suggest that Late Jurassic–Early Cretaceous porphyry Cu deposits in this region were formed in an island arc setting above a northeast-directed Neotethyan subduction zone (the Somkheto– Karabakh island arc), although an accreted arc margin seems more likely from Fig. 2. 4.2.3. Iran Mesozoic rocks related to subduction of Neotethyan ocean basins are preserved in the Sanandaj–Sirjan Zone (which runs NW–SE across central Iran; Fig. 3), the eastern Alborz–Kopeh Dagh (Sabzevar zone, NE Iran), and the Central Iranian microcontinent (eastern Central Iran). There is considerable debate regarding the detailed history of Neotethyan basin opening and closure in Iran, but evidence from several ophiolite belts across the country strongly suggests that several small (back-arc?) basins existed in addition to the main Neotethyan ocean, which was finally eliminated by collision between Iran and Arabia along the Zagros suture zone in the Neogene. The main Neotethyan ocean opened when the Sanandaj–Sirjan Zone and Central Iranian microcontinent (Lut and Tabas blocks) rifted from the northeastern margin of Gondwana in the Early Triassic (Hooper et al., 1994; Fig. 1). By the Late Triassic or Early Jurassic these Cimmerian continental fragments had collided with the southern margin of Eurasia, and northward-directed subduction of the Neotethys ocean began beneath this accreted margin (Fig. 2A; Horton et al., 2008; Masoodi et al., 2013; Mirnejad et al., 2013a). Late Triassic–Cretaceous I-type arc plutons relating to this period of subduction occur throughout the Iraqi Zagros Suture Zone, the Sanandaj–Sirjan Zone, and the Makran (Agard et al., 2005; Ahmadi Khalaji et al., 2007; Ali et al., 2013; Aliani et al., 2012; Arvin et al., 2007; Azizi and Jahangiri, 2008; Mahmoudi et al., 2011; Mohajjel and Fergusson, 2014; Shahabpour, 2010; Shahbazi et al., 2010). However, as in the Pontides of Turkey, no significant porphyry deposits of this age are preserved, likely due to erosion down to batholithic levels. Small orogenic gold deposits of Late Cretaceous–Paleogene age have been reported in the Sanandaj–Sirjan Zone, analogous to other convergent margin mesothermal gold deposits worldwide (Aliyari et al., 2012). Back-arc rifting along the accreted margin opened several small ocean basins such as the Sabzevar and Sistan oceans in northeast and eastern Iran, respectively (Fig. 2B). Closure of these basins in the Late Cretaceous–Oligocene gave rise to calc-alkaline arc magmatism in the Arghash Massif (Alaminia et al., 2013; Jafari et al., 2013) and Lut block
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(Arjmandzadeh et al., 2011; Kaz'min and Tikhonova, 2008). However, in a situation analogous to the highly oblique subduction of the Vardar ocean in the Balkans, the polarity of subduction of the Sistan ocean is disputed (Bröcker et al., 2013; Saccani et al., 2010; Tirrul et al., 1983). Considerable debate surrounds the history of collision(s) between the Sanandaj–Sirjan Zone, Central Iran, and Arabia. The paleogeographic reconstructions in Fig. 2 indicate that the Sanandaj–Sirjan Zone docked with Eurasia in the Mesozoic (although a small back-arc basin may have opened between these blocks in the Middle Cretaceous; Ghasemi and Talbot, 2006), and suggest that collision with Arabia in the Neogene was the last phase of Neotethyan closure. However, several authors contend that the Sanandaj–Sirjan Zone collided first with Arabia in the Cretaceous (Alavi, 1980, 1994, 2004) or Oligocene (Hooper et al., 1994), prior to collision with Central Iran in the late Cenozoic (e.g., Ghalamghash et al., 2009; Glennie, 2000). Part of the confusion may relate to the shift of arc magmatism from the Sanandaj–Sirjan Zone in the Mesozoic to the Urumieh–Dokhtar magmatic belt in Central Iran in the Late Cretaceous–Paleogene. Glennie (2000) relates this switch to the closure of two separate Neotethyan basins to the south and north of the Sanandaj–Sirjan Zone. However, an alternative explanation may be that the angle of Neotethyan subduction shallowed in the Cretaceous, leading to a shift in the axis of arc magmatism from the Sanandaj–Sirjan Zone to the parallel but more northeasterly Urumieh–Dokhtar belt in the Paleogene (Fig. 3; Verdel et al., 2011; Mohajjel and Fergusson, 2014). Ophiolitic fragments found between the Sanandaj–Sirjan Zone and Urumieh–Dokhtar belt may relate to closure of a small back-arc basin (e.g., Ghasemi and Talbot, 2006), but there does not seem to be time or space to open up a large ocean basin between these blocks in the Cretaceous (see Fig. 2). Collision with small island arc terranes (rather than the Arabian continent) may explain the evidence of moderate collisional deformation on the southern margin of the Sanandaj–Sirjan Zone in the Late Jurassic– Early Cretaceous (Azizi and Asahara, 2013). 4.2.4. Makran (Pakistan) Late Cretaceous–late Paleocene basaltic andesitic to andesitic volcanism in the Chagai region of southwestern Pakistan (Fig. 3) is partly submarine and island arc in character (Arthurton et al., 1982; Nicholson et al., 2010; Richards et al., 2012; Siddiqui, 2004). Although this is not evident in the paleogeographic reconstruction shown in Fig. 2A, it appears that the Chagai arc was initially oceanic, and collided with the Afghan (Helmand) block in the Late Cretaceous, whereafter the arc changed to continental in character (Nicholson et al., 2010). No mineralization is known to be associated with the Mesozoic arc history of this belt, but major porphyry deposits formed in the Cenozoic continental arc (see Sections 4.3.3 and 4.4.3). 4.3. Paleogene 4.3.1. Turkey The main events that affected Turkey in the Paleogene were subduction of Neotethyan oceanic lithosphere attached to the African plate along the Crete and Cyprus trenches, and collision of the Tauride–Anatolide block (TAB) with the Pontide and Kırşehir blocks in the late Paleocene–early Eocene to eliminate the Ankara–Erzincan branch of the northern Neotethys (Fig. 2C; Aldanmaz et al., 2000; Ilbeyli et al., 2004; Boztuğ and Jonckheere, 2007; Kaymakci et al., 2009; Arslan et al., 2013). Tectonic readjustments following this collisional event in the Eocene caused inversion in the Pontides (Espurt et al., 2014), and triggered a major flare-up of calc-alkaline magmatism across northern Turkey in response to slab rollback and breakoff (Boztuğ and Arehart, 2007; Boztuğ and Harlavan, 2008; Boztuğ et al., 2006; Kaygusuz et al., 2008; Kaymakci et al., 2010; Keskin et al., 2008; Önal et al., 2005). Postcollisional extensional tectonics and upwelling of asthenospheric mantle caused partial melting of previously subduction-modified
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lithosphere throughout Anatolia (Altunkaynak, 2007; Altunkaynak and Dilek, 2013; Sarıfakıoğlu et al., 2013), whereas delamination of subcontinental lithospheric mantle (SCLM) beneath the Pontides caused crustal melting (Arslan and Aslan, 2006; Arslan et al., 2013; Aslan et al., 2014; Aydin et al., 2008; Karsli et al., 2010; Temizel et al., 2012; Topuz et al., 2011). Several small porphyry and epithermal deposits are associated with Paleocene–Eocene post-collisional calc-alkaline magmatism in the Pontides, but no economic deposits have been discovered to date (Akçay and Gündüz, 2004; Yigit, 2009). To the north of the Pontides, back-arc extensional tectonics are thought to have widened the Black Sea during the early Paleogene (Shillington et al., 2008), although controversy remains over the history of this basin (e.g., Arslan et al., 2013; Eyuboglu et al., 2011, 2012). In southern Turkey, northward subduction of the main Neotethyan ocean continued along the Bitlis–Zagros subduction zone (Fig. 3), giving rise to Late Cretaceous–Eocene magmatism in the Maden-Helete arc of SE Anatolia, and back-arc magmatism along ENE-trending strike-slip faults in the Central and Eastern Taurides (Robertson et al., 2007). Distinguishing between a back-arc and collisional origin for late Paleogene calc-alkaline magmas in the TAB is difficult but not unimportant, because this magmatism is associated with major porphyry Cu–Au and epithermal Au mineralization in the Çöpler–Kabataş district (~44 Ma; Imer et al., 2013; Kuşcu et al., 2013), and Oligocene porphyry Cu–Au–Mo deposits in the Tunceli region (~25 Ma; Imer et al., in press) (Figs. 2C and 3). Deposits of slightly younger age but equivalent geodynamic setting occur to the southeast along the orogen in Iran (see Section 4.3.2). From the Oligocene through the Neogene and Quaternary, magmatism throughout central and eastern Turkey became increasingly localized and alkaline in character, reflecting final closure of the Bitlis–Zagros subduction zone in the Miocene, and a transition to fully post-collisional tectonics (Akay, 2009; Boztuğ and Jonckheere, 2007; Yılmaz et al., 2001). In contrast, extensional tectonics in the Aegean and western Anatolia since the late Oligocene are related to slab rollback or differential upper plate advance on the Hellenic trench (Agostini et al., 2010). This distinct tectonic setting gave rise to structurally localized calc-alkaline magmatism and significant porphyry Cu–Mo–Au and high-sulfidation epithermal Au mineralization in the Biga Peninsula (e.g., the TV Tower epithermal Au deposit, and the Halilaga Cu–Au and Tepeoba porphyry Cu–Mo–Au deposits; Yigit, 2012). This mineralization is thought to be correlative with the Serbo-Macedonian–Rhodope metallogenic belt in the Balkans. 4.3.2. Iran Arc magmatism in the Sanandaj–Sirjan zone largely ceased at the end of the Cretaceous, and shifted to the Urumieh–Dokhtar magmatic belt, approximately 100 km to the northeast (Fig. 3). As noted above, researchers disagree about the relationship between the Sanandaj– Sirjan and Urumieh–Dokhtar arcs, but the simplest explanation would appear to be that the angle of subduction flattened, and the NW–SE axis of magmatism shifted inland to the northeast, to form a new arc (the Urumieh–Dokhtar arc) in the Paleogene (Mohajjel and Fergusson, 2014; Whitechurch et al., 2013). Verdel et al. (2011) have suggested that the Paleocene–Eocene magmatic flare-up in the Urumieh–Dokhtar arcs occurred as a previously flat-subducting slab rolled back prior to Miocene collision, exposing hydrated lithosphere to an influx of hot asthenospheric material in the re-opened mantle wedge. Alternatively, Ahmadian et al. (2009) and Allen (2009) have suggested that middle Eocene magmatism in the Urumieh–Dokhtar belt and in Central Iran was related to back-arc extension in advance of final collision. A small back-arc basin between the Sanandaj–Sirjan zone and Urumieh– Dokhtar belt may also have closed at this time (Ghasemi and Talbot, 2006). Cenozoic magmatism in the Urumieh–Dokhtar belt has been the focus of considerable study, because it is related to several large
porphyry Cu deposits, mainly of Miocene age. Eocene magmatism was initially calc-alkaline in character, but changed to more potassic (shoshonitic) compositions in the Oligocene and early-middle Miocene, and then to even more alkaline compositions in the late Neogene (Hassanzadeh, 1993). This transition is believed to reflect the onset of final collision between Arabia and the accreted Eurasian margin, whose timing has been estimated to be anywhere between Late Cretaceous to mid-Miocene. A Miocene age for final collision is preferred by most recent authors, and is consistent with the Miocene age of collision along the Bitlis subduction zone in Turkey. However, the irregular shape of the Arabian indentor suggests that collision was likely diachronous along the belt (Agard et al., 2005), which may account for some of the variability in estimates for the timing of collision. It also seems likely that the collision was initially relatively “soft” (Ballato et al., 2011), perhaps beginning in the Eocene (Agard et al., 2005; Allen and Armstrong, 2008; Dargahi et al., 2010; Hafkenscheid et al., 2006; Horton et al., 2008; Mazhari et al., 2009), with final “hard” collision occurring in the late Oligocene (McQuarrie and van Hinsbergen, 2013) or early-middle Miocene (Ali et al., 2013; Allen et al., 2004; Ballato et al., 2011; Förster, 1978; Karagaranbafghi et al., 2012; Mohajjel and Fergusson, 2014; Stoneley, 1981). In particular, Mouthereau et al. (2012) suggest that collision began in the late Eocene at ~35 Ma, and was followed by crustal thickening in the Oligocene, and uplift of the central Iranian plateau in the middle Miocene (15–12 Ma) prior to slab breakoff. This timing is consistent with thrust-driven uplift of the High Zagros (Arabian margin), which had developed by the Oligocene or early Miocene (Fakhari et al., 2008; Gavillot et al., 2010). Behind the main Neotethyan destructive margin, Eocene volcanism also occurred in the Alborz of north and northeastern Iran, as well as the Lut block in eastern Iran. Although the detailed tectonic history of this area is not well known, it appears to have involved the formation and destruction of one or more small back-arc ocean basins, such as the Sabzevar and Sistan oceans. Calc-alkaline Eocene magmatism has thus variably been interpreted to be of subduction (Alaminia et al., 2013; Arjmandzadeh and Santos, 2014; Richards et al., 2012; Spies et al., 1984), back-arc (Asiabanha and Foden, 2012), or post-collisional origin (Castro et al., 2013; Nabatian et al., 2014; Pang et al., 2013b). Several small porphyry deposits occur in association with these Eocene sequences, including the Shadan and Maherabad prospects in the Lut block (Figs. 2C and 3; Richards et al., 2012; Siahcheshm et al., 2012, 2014), and a few small intermediate-sulfidation epithermal gold deposits are reported in the Alborz (Shamanian et al., 2004). By the Oligocene, magmatism in northeastern Iran had a clear collisional character (Asiabanha et al., 2012), and oroclinal bending and uplift had begun in the Kopeh Dagh (Hollingsworth et al., 2010). A similar transition is recognized in northwest Iran, where calcalkaline magmatism in the early Oligocene gives way to shoshonitic plutonism in the late Oligocene (Aghazadeh et al., 2011). 4.3.3. Makran (Pakistan) The Chagai island arc appears to have accreted to the Afghan (Helmand) block by the Eocene–Oligocene, and subsequent Oligocene– Quaternary magmatism is of continental arc affinity (Siddiqui, 2004). The axis of the arc remained fixed, however, such that successive sequences or volcanic rocks overlie each other, and intrusions overlap. This spatial coincidence may explain the progressive increase in size of porphyry Cu–Au deposits as the arc evolved, with no known deposits associated with Cretaceous–Paleocene rocks, only small prospects in the middle-late Eocene (Ziarate; Perelló et al., 2008), but several large deposits such as Saindak and Reko Diq occurring in the Miocene (Perelló et al., 2008; Richards et al., 2012; Sillitoe, 1979; Sillitoe and Khan, 1977). The Makran accretionary complex developed to the south of the Chagai arc over a shallow-dipping subduction zone since the Late Cretaceous (Jacob and Quittmeyer, 1979), and is one of the largest sedimentary prisms on Earth (McCall, 1997; White, 1979). In the late
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Oligocene or early Miocene, fold-and-thrust deformation began in this belt, and dextral strike-slip motion was initiated on the Chaman transform zone to the east (Fig. 3) as the effects of the Indian plate collision began to be felt along the eastern Eurasian margin (Lawrence et al., 1981). 4.4. Neogene–Quaternary 4.4.1. Turkey Final collision of Arabia with Turkey along the Bitlis–Zagros subduction zone likely occurred in the early-middle Miocene (Fig. 2D; Şengör and Yilmaz, 1981; Pearce et al., 1990; Yilmaz, 1993; Robertson et al., 2007; Hüsing et al., 2009; Çolakoğlu and Arehart, 2010; Okay et al., 2010). This was followed by crustal thickening of the Turkish–Iranian plateau, and subsequent westward strike-slip escape of Turkey (Allen, 2010; Allen et al., 2004). Uplift in Anatolia has been interpreted to be due to delamination of subcontinental mantle lithosphere (Göğüs and Pysklywec, 2008; Sandvol et al., 2003), and/ or slab breakoff and asthenospheric upwelling (Lei and Zhao, 2007; Schildgen et al., 2014; Şengör et al., 2008; Zor, 2008). The result has been to thin the lithosphere in this region (to 60–80 km) while thickening the crust (to 30–55 km), and to trigger widespread postcollisional volcanism (Angus et al., 2006). Oroclinal bending in the eastern Pontides and Lesser Caucasus at this time (late Miocene– Pliocene) was caused by the irregular shape of the Arabian indentor (Hisarlı, 2011). Post-collisional magmatism ranged from calc-alkaline to shoshonitic in the early-middle Miocene, to progressively more alkaline compositions in the late Miocene to Quaternary (Aldanmaz et al., 2000; Altunkaynak and Dilek, 2006; Altunkaynak and Genç, 2008). These younger alkaline volcanic centers are commonly localized by major extensional or transtensional structures, reflecting E–W tectonic adjustments in response to N–S collision. They are found throughout Turkey, including along the North Anatolian Fault Zone (Temel et al., 2010), northwest Anatolia (Aldanmaz, 2002; Ersoy et al., 2010, 2012; Yılmaz et al., 2001), southwest Anatolia (Prelevic et al., 2012), central Anatolia (Kuşcu and Geneli, 2010; Ocakoğlu, 2004; Şen et al., 2004), eastern Anatolia (Karsli et al., 2008; Kürüm et al., 2008; Özdemir et al., 2006; Pearce et al., 1990; Yilmaz et al., 1998), the eastern Pontides (Aydin et al., 2008; Keskin, 2003; Kheirkhah et al., 2009), and northeast Syria (Lease and Abdel-Rahman, 2008). In most cases, the geochemical and isotopic compositions of these magmas indicate derivation from previously subduction-modified lithosphere, with variable contributions from asthenospheric melts. Melting is generally attributed to slab rollback and breakoff (op. cit., and Reilinger et al., 2006; Le Pichon and Kreemer, 2010; Karaoğlu and Helvacı, 2014). Few of these post-collisional volcanic centers in central and eastern Turkey are known to be associated with mineralization, although it might be speculated that unexposed alkalic-type porphyry and epithermal deposits could exist below the current shallow levels of exposure (e.g., Richards, 1995, 2009). In contrast, in the extensional tectonic environment of western Anatolia, Miocene calc-alkaline volcanic complexes are associated with several major low-sulfidation (Ovacik), intermediate-sulfidation (Efemçukuru, Kucukdere), and high-sulfidation porphyry-related (Kisladag) Au deposits (Figs. 2D and 3; Yilmaz, 2003; Yilmaz et al., 2007; J.W. Hedenquist, personal communication, 2014). These volcanic systems and gold deposits developed in an E–W extensional environment related to Aegean tectonics, as noted above. 4.4.2. Iran The controversy over the timing of collision between Arabia and Eurasia along the Zagros section of the Bitlis–Zagros subduction zone has been discussed above. The most likely timing seems to be in the Miocene, perhaps propagating diachronously from early Miocene in the northwest of Iran, to late Miocene in the southeast (Chiu et al.,
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2013; Hafkenscheid et al., 2006; Hooper et al., 1994; Robertson et al., 2009). Neogene dextral strike-slip faulting lengthened the orogen until ~5 Ma, when eastward extension was blocked by collision of the Afghan (Helmand) plate with India (Allen et al., 2011). Allen et al. (2013b) suggest that crustal thickening has occurred by shortening across the orogen (~68 km across the Zagros, and ~120 km across the Arabian plate; see also Regard et al., 2005). Slab breakoff following collision is widely considered to have caused uplift and post-collisional magmatism throughout Iran, although the timing of breakoff is debated (mid-late Miocene: van Hunen and Allen, 2011; Mouthereau et al., 2012; Plio-Pleistocene: Snyder and Baranzangi, 1986; Molinaro et al., 2005). Whether or not there has been delamination of the SCLM below the Zagros orogen is also debated (yes: Shomali et al., 2011; no: Paul et al., 2010). Neogene–Quaternary mafic post-collisional volcanism across Iran (from the northwest to the Lut block) is potassic to ultrapotassic in character, and is generally thought to have a subduction-modified mantle lithosphere origin (Ahmadzadeh et al., 2010; Allen et al., 2013a; Kheirkhah et al., 2013; Pang et al., 2013a), perhaps with an asthenospheric component in the case of the isolated Damavand stratovolcano (Davidson et al., 2004; Liotard et al., 2008; Mirnejad et al., 2010; Shabanian et al., 2012). In many cases, the volcanism is localized along transtensional structures (e.g., Ahmadzadeh et al., 2010). Pang et al. (2012) have proposed that alkali basaltic volcanism associated with N–S strike-slip faults in eastern Iran (Neh and Nayband faults, Fig. 3; Meyer and Le Dortz, 2007) formed in response to extension following delamination of thickened SCLM (a delayed response to Late Cretaceous collision between the Lut and Afghan blocks). In the Urumieh–Dokhtar magmatic belt, bimodal felsic and mafic alkalic magmatism occurred in the Miocene, following earlier calc-alkaline (Eocene) and shoshonitic (Oligocene) magmatism (Hassanzadeh, 1993). Omrani et al. (2008) have suggested slab breakoff as a cause for adakite-like Pliocene–Quaternary magmatism in this belt, and Shafiei et al. (2009) and Shafiei (2010) have proposed that Miocene post-collisional calc-alkaline magmas involved a significant fraction of remobilized subduction-modified lower crust in their formation. Shafiei et al. (2009) in particular argue that this process is key to the fertility of these mid-Miocene magmas, which host most of the largest porphyry Cu–Mo deposits in Iran (e.g., Sar Cheshmeh, Meiduk; Figs. 2D and 3; McInnes et al., 2003; Taghipour et al., 2008; Mirnejad et al., 2013b). Mirnejad et al. (2013b) report an age range of 15.1–9.8 Ma (Re–Os on molybdenite) for porphyry deposits in the Kerman section of the Urumieh–Dokhtar belt, which brackets the ages of Sar Cheshmeh (13.6 ± 0.1 Ma, zircon U–Pb; McInnes et al., 2003) and Meiduk (12.5 ± 0.1 Ma, zircon U–Pb, McInnes et al., 2003; 12.23 ± 0.07 Ma, molybdenite Re–Os, Taghipour et al., 2008). Several smaller porphyry Cu (e.g., Darreh-Zerreshk and Ali-Abad; ~ 16 Ma; Zarasvandi et al., 2005, 2007) and porphyry Au deposits (e.g., Dalli; ~20 Ma; Ayati et al., 2013) formed in the early Miocene in Central Iran, and the large Sungun porphyry Cu–Mo deposit in NW Iran (Fig. 3) formed at 21.1–19.5 Ma (Aghazadeh et al., 2012). All of these deposits are variably attributed to post-collisional magmatic processes (e.g., Jamali et al., 2010), and the increasing ages along the orogen to the northwest are consistent with the diachronous collision of Arabia noted above. Epithermal Au mineralization also formed in the middle-late Miocene in NW Iran, including the Sari Gunay alkalic-type Au deposit (Fig. 3; ~ 10.7 Ma; Richards et al., 2006), and the Zarshuran (14.2 ± 0.4 Ma; Mehrabi et al., 1999) and Agdarreh sediment-hosted Au deposits (Asadi et al., 2000; Daliran, 2008). Oroclinal bending (analogous to the Caucasus) began in the Kopeh Dagh of NE Iran in the early-to-middle Oligocene (Hollingsworth et al., 2010), followed by uplift and shortening in both the Alborz and Kopeh Dagh in the Miocene (Allen et al., 2003; Guest et al., 2006). These mountain belts effectively mark the northern limit of deformation in the Eurasian margin caused by the Arabian collision. The region has
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subsequently been affected by transpressional tectonics with strike-slip faulting in the Pliocene–Quaternary (Bachmanov et al., 2004; Bonini et al., 2003; Hollingsworth et al., 2006; Shabanian et al., 2010). In the Makran of southeast Iran, subduction of the Neotethys continues, as reflected by the recently active Bazman and Taftan volcanoes, which both tap an arc source. The Bazman volcano is located very close to the margin between continental collision and normal subduction, and Quaternary basalts erupted from parasitic cones around this volcano may be derived from subduction-modified lithospheric mantle (Saadat and Stern, 2011). The more easterly Taftan volcano shows a stronger arc signature (Biabangard and Moradian, 2008), and is likely correlative with the Koh-i-Sultan volcano in the Chagai arc of Pakistan (Jacob and Quittmeyer, 1979). A small low-sulfidation epithermal Au deposit occurs near Bazman (Chahnali; Daliran et al., 2005), and subeconomic (but minimally explored) porphyry and highsulfidation-style epithermal mineralization occurs around the Taftan volcano. 4.4.3. Makran (Pakistan) Calc-alkaline to mildly alkaline volcanic rocks were erupted in the Chagai arc from the Miocene to Quaternary, with recent volcanism being recorded in the large Koh-i-Sultan stratovolcano (Richards et al., 2012; Siddiqui, 2004). The arc extends for ~450 km from the Bazman volcano in Iran to the eastern end of the Chagai belt. Siddiqui (2004), Perelló et al. (2008), and Richards et al. (2012) describe the evolution of this arc, and the formation of a series of porphyry Cu–Au–Mo deposits of increasing size from the mid-Eocene to the mid-Miocene, culminating in the giant Reko Diq cluster of porphyries at ~ 11 Ma (Fig. 3; Perelló et al., 2008; Razique et al., 2014). Reko Diq is one of the world's largest unmined Cu–Au deposits (resources of 5.9 Gt @ 0.41% Cu and 0.22 g/t Au1), while the nearby Saindak porphyry Cu–Mo deposit (22.30 ± 0.05 Ma; Richards et al., 2012) is currently being mined by Metallurgical Corporation of China Ltd., with an estimated resource of 440 Mt @ 0.41% Cu, 0.002% Mo (Singer et al., 2008) plus a significant amount of Au. 5. Pamir–Himalayas–Indochina The Himalayan orogenic belt has been studied in a somewhat fragmental way from the perspectives of Pakistan, India, Tibet, and Indochina, such that it can sometimes be difficult to correlate across these national borders. Political and security concerns in northwestern Pakistan, Kashmir and Tibet have further limited the extent of modern geological research, although there has been a recent surge of publications relating to Tibet and its porphyry deposits. In reviewing the tectonics and metallogeny of this belt, I have drawn from several key overview papers (Beaudoin et al., 2005; Bierlein et al., 2009; Chatterjee et al., 2013; Dewey et al., 1988; Metcalfe, 2013; Pirajno, 2013; Qin et al., 2012; Yin and Harrison, 2000), as well as individually cited papers. Although the Pamir–Himalayas–Indochina segment of the Tethyan orogen is unified by the effects of the Paleogene Indian continental collision, the belt is heterogeneous, and its pre-collisional history is more easily described in terms of two contiguous sections, Kohistan–Pamir and Tibet–Indochina, during the Mesozoic. 5.1. Paleotethys As for most of the Tethyan orogen, the history of the Paleotethys ocean is not well known in the Himalayan region and Indochina. Dewey et al. (1988), Yin and Harrison (2000), Metcalfe (2006, 2013), and T.N. Yang et al. (2014) describe a complex assemblage of terranes in southern and southeast Asia, which were initially rifted from 1 Data from Tethyan Copper Company Ltd. (http://www.tethyan.com/ TheRekoDiqProject/RekoDiqResource.aspx; accessed 10 February, 2014).
Gondwana in the Paleozoic (Zhu et al., 2011b), and then accreted progressively to the Asian margin from the Paleozoic to Cenozoic by closure of multiple Tethyan ocean basins. Related to these ocean closure events, Schwab et al. (2004) describe several periods of arc magmatism in the Pamir region, beginning in the early Paleozoic (~ 575–410 Ma) and continuing until the Triassic in the north-facing Kunlun arc. Similarly, Hu et al. (2013) report the occurrence of Cambrian arc volcanism in the Lhasa terrane of southern Tibet, and Guo et al. (2012) describe Middle Triassic arc magmatism (244–234 Ma) related to northward Paleotethyan subduction along the northeastern margin of Tibet. The latter magmatism may correlate with Triassic volcanism in the Kunlun arc to the west, and the Mianlue arc in the Qinling terrane further east. No porphyry deposits are known to be associated with Paleozoic sequences in this section of the Tethyan orogen, although numerous deposits of Paleozoic age occur in the Central Asian orogenic belt to the north (Fig. 3), including the giant Oyu Tolgoi porphyry Cu–Au– (Mo) deposit (372 Ma; Perelló et al., 2001; Crane and Kavalieris, 2012). Tibet consists of several terranes of Gondwanan affinity that accreted to the Asian margin from the late Permian through the Mesozoic through closure of multiple Paleotethyan and Mesotethyan ocean basins (Metcalfe, 2006). The most northerly of these is the SongpanGanzi terrane (Fig. 3), which was accreted to Kunlun (Tarim–North China) in the Late Permian. The Qiangtang terrane (which has an internal suture, suggesting that it is itself a compound terrane; Zhao et al., 2014) was then accreted to Songpan-Ganzi in the Late Triassic along the Jinshajiang suture (Fig. 3). The Lhasa terrane then collided with Qiangtang along the Bangong-Nujiang suture (Fig. 3) in the Late Jurassic (Dewey et al., 1988; Yin and Harrison, 2000) or Late Cretaceous (Liu et al., 2013; Y. Li et al., 2013; Zhu et al., 2011a) (Figs. 2A, B, and 3). Early Permian–Early Triassic (275–248 Ma) calc-alkaline volcanic rocks in the Qiangtang terrane (Wang et al., 2008a; Yang et al., 2011) and Triassic–Jurassic (225–193 Ma) calc-alkaline plutons in the HohXil–Songpan-Ganzi complex (Zhang et al., 2014) relate to closure of at least two ocean basins prior to or during suturing of the Lhasa terrane (Roger et al., 2010). Although these older arc sequences are not widely preserved, a number of Jurassic and Cretaceous porphyry deposits have been found in these terranes. The largest of these occur in the Middle Jurassic Xiongcun porphyry Cu–Au district in the southern Lhasa terrane (which includes the Xietongmen and Newtongmen deposits; ~ 174 Ma; Tafti et al., 2009; Lang et al., 2014), and the Early Cretaceous Bangongco metallogenic belt in the southern Qiangtang terrane (which includes the Duolong porphyry Cu–Au deposit (~ 115 Ma; J. Li et al., 2011 J. Li et al., 2012; J.-X. Li et al., 2013)) (Fig. 3). These deposits are thought to be normal arc porphyry systems, related to subduction of Mesotethys oceanic lithosphere. Early Mesozoic porphyry deposits also occur in the Indosinian porphyry belt of SE China, which formed in the Triassic Zhongdian island arc (see T.N. Yang et al., 2014, for a detailed description of the geological history of this complex late Paleozoic to early Mesozoic collisional belt). These include the large Pulang (~213 Ma; W. Li et al., 2011) and Yangla porphyry Cu and skarn deposits (~233 Ma; X.-A. Yang et al., 2014; Zhu et al., in press) (Fig. 3), and are suggested to have formed in response to westward subduction of the Ganzi-Litang and Jinshajiang ocean basins, respectively, beneath the Yidun terrane (Wang et al., 2011; Zhu et al., in press).
5.2. Late Mesozoic 5.2.1. Kohistan–Pamir The Kohistan arc of northern Pakistan is relatively well studied (e.g., Bignold et al., 2006; Coward et al., 1982, 1987; Jagoutz and Schmidt, 2012; Jagoutz et al., 2006, 2007, 2009; Khan et al., 1989,
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1993; Petterson, 2010; Searle et al., 1999), although much less is known about the Mesozoic history of the Pamir. The Kohistan terrane is considered to have formed as a Cretaceous (~134–90 Ma) intra-oceanic arc overlying a north-directed subduction zone (Fig. 2A, B; Bard, 1983; Bignold and Treloar, 2003; Petterson and Treloar, 2004; Petterson, 2010; Bouilhol et al., 2011). It is of significant geological interest because subsequent collisional uplift has exposed an almost complete section of oceanic arc crust, from upper crustal volcanic sequences (the Chalt Volcanic Group) and coeval tonalitic intrusions of the early Kohistan batholith, through gabbroic intrusions and layered mafic cumulates of the lower crustal Chilas Complex, to the lowermost mafic–ultramafic Kamila amphibolites and Jijal Complex (Bignold et al., 2006). The petrology of these unique lower crustal rocks has been studied in detail by Jagoutz and co-workers (op. cit.) and Dhuime et al. (2009). While the origin of the Kohistan arc is relatively well established, debate arises concerning its collisional history, with some authors proposing accretion first to the Indian plate margin in the Late Cretaceous at ~ 85 Ma or later (Bard, 1983; Bouilhol et al., 2013; Chatterjee et al., 2013; Khan et al., 2009), and others arguing for initial accretion to the Asian margin at ~ 100–75 Ma (Bignold et al., 2006; Petterson, 2010; Petterson and Treloar, 2004; Rehman et al., 2011; Searle et al., 1999; Treloar et al., 1996). Whichever interpretation is correct, the collision caused Late Cretaceous crustal thickening, amphibolite to granulite facies metamorphism of the lowermost crustal sequences (Bouilhol et al., 2011; Jan and Howie, 1981; Petterson, 2010; Petterson and Treloar, 2004), and partial melting to form calc-alkaline gabbro– granodiorite plutons (85–40 Ma; Bignold et al., 2006). Despite preservation of upper crustal volcanic and subvolcanic sequences, no porphyry-type deposits are known in Kohistan. However, small deposits of Cu–Au–PGE-bearing sulfides occur in dunites in the Chilas Complex (author's unpublished data), and chromite is mined from layers or pods in the ultramafic Jijal Complex. 5.2.2. Tibet–Indochina The Cretaceous period saw widespread arc magmatism across the Tibetan plateau, with uplift beginning as an Andean-type margin in the Late Cretaceous (~ 85 Ma), and reaching its maximum extent by ~ 45 Ma (Ding et al., 2014; Rohrmann et al., 2012). Early Cretaceous adakite-like calc-alkaline intrusions in the northern Lhasa terrane have been related to slab breakoff in the south-directed Bangong– Nujiang subduction zone (H. Wu et al., in press), whereas similar Early Cretaceous intrusions in the southern Gangdese magmatic belt of the Lhasa terrane are interpreted to have been generated by melting of Neotethyan oceanic lithosphere (Zhu et al., 2009). In contrast, Late Cretaceous adakite-like intrusions in the Gangdese belt are interpreted to have been derived by partial melting of juvenile mafic lower crust, underplated during Cretaceous subduction, and thickened by flat subduction from ~ 80–70 Ma (Wen et al., 2008a,b). In SE China (close to Pulang), a late- or post-collisional origin (with respect to collision between the Lhasa and Qiangtang terranes) has been proposed for Late Cretaceous adakite-like intrusions and associated Mo–Cu mineralization (X.-S. Wang et al., 2014). 5.3. Paleogene It is now generally agreed that India began to collide with the accretionary Asian margin in the early Eocene at ~ 55–50 Ma (Guillot et al., 2003; Khan et al., 2009; Meng et al., 2012; Rehman et al., 2011; Searle et al., 1999; Treloar et al., 1996; Zhang et al., 2012), although some authors call for an earlier onset (e.g., ~ 70–65 Ma; Cai et al., 2011), and others for a later event (e.g., ~40 Ma; Bouilhol et al., 2013; ~ 34 Ma; Aitchison et al., 2007; Fig. 2C, D). Some of the disagreement in these estimates may arise because collision was likely not a single event, and may have been diachronous along the orogen (White and Lister, 2012). The Neotethyan oceanic lithosphere slab is thought to
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have broken off at ~45 Ma (Hafkenscheid et al., 2006; Replumaz et al., 2010; van der Voo et al., 1999). Initial collision of the Indian plate is thought to have involved a wide passive continental margin known as Greater India (Ali and Aitchison, 2005; van Hinsbergen et al., 2012), which subducted below Tibet until final contact with the Indian craton in the late Eocene (~34 Ma; Meng et al., 2012). Upper crustal shelf sediment sequences were scraped off the Greater India plate and accreted to form the Himalayas, while the underlying dense lower crust and mantle lithosphere were subducted (Capitanio et al., 2010). This early Eocene collisional event has been termed a “soft collision”, in advance of the later “hard collision” with the main Indian continental mass (Fig. 2C, D; Meng et al., 2012; van Hinsbergen et al., 2012). The Greater India collision resulted in slowing of the India–Asia convergence rate from ~ 18 cm/yr (Capitanio et al., 2010) to ~ 4.7 cm/yr (Guillot et al., 2003). Rapid convergence in the Late Cretaceous may have been caused by impact of the Morondova mantle plume (which split Madagascar from India at this time), and was sustained through the Paleocene–early Eocene by increased slab-pull and ridge-push forces (van Hinsbergen et al., 2011). I-type calc-alkaline arc magmatism continued during the run-up to the initial collision, with Paleocene volcanosedimentary sequences forming in central and northern Kohistan (Dir Group and Shamran volcanics; Sullivan et al., 1993), widespread Paleocene–early Eocene Linzizong volcanic sequences being erupted across southern Tibet (Coulon et al., 1986; Mo et al., 2008), and Paleocene–early Eocene gabbroic–granitic intrusions being emplaced along the length of the orogen to form the Kohistan, Ladakh, and Gangdese batholiths (Bignold et al., 2006; White et al., 2011; Zhang et al., 2013). In detail, the Ladakh batholith appears to have formed episodically between 68–45 Ma, with the bulk being emplaced between 63–55 Ma (White et al., 2011). Despite the good preservation and exposure of these shallow crustal volcanic sequences and subvolcanic plutons, few porphyry-type deposits are known from this period in the central and western parts of Tibet; only two small porphyry Cu (Jiru; ~ 45 Ma; Zheng et al., 2014) and porphyry Mo (Sharang; ~52 Ma; Zhao et al., 2012, 2013) deposits are reported in the literature. R. Wang et al. (2014b,c) have suggested that relatively low magmatic water contents and oxidation states in these Paleogene magmas, resulting from late-stage subduction processes during the onset of collision, may explain their low metallogenic potential. In comparison, the early-middle Eocene Yulong porphyry belt in the Qiangtang terrane of eastern Tibet (Fig. 3) hosts several large Cu–Au deposits, including Yulong (40.1 ± 1.8 Ma), Machangqing (35.8 ± 1.6 Ma), Habo (35.5 ± 0.2 Ma; Zhu et al., 2013), and Xifanping (32.1 ± 1.6 Ma) (ages from Hou et al., 2003, 2006, except where otherwise indicated). Associated potassic magmatism with adakite-like trace-element characteristics is interpreted to have formed by partial melting of previously subduction-modified lithospheric mantle (Jiang et al., 2006) or juvenile lower arc crust (G. Li et al., 2012; Lu et al., 2013a, 2013b), and to have been emplaced in pull-apart basins along collision-related NNW–SSE dextral strike-slip faults (Hou et al., 2003; Lu et al., 2013b). Compressional forces between ~ 45–30 Ma in the upper Tibetan plate caused crustal thickening (Chung et al., 2009; Searle et al., 2011), possible delamination or erosion of the Tibetan SCLM (Chung et al., 2009), and uplift of the Tibetan plateau to its maximum levels by ~45 Ma (Rohrmann et al., 2012). These tectonic changes are reflected in the character of magmatism, which became distinctly more alkaline in the late Eocene. Potassic–ultrapotassic lavas and adakite-like mafic– intermediate and granitoid plutons formed at this time (46–30 Ma), and are generally thought to reflect partial melting of thickened Tibetan mafic lower crust (Chen et al., 2013; Chung et al., 2009; Searle et al., 2011; Wang et al., 2008b), possibly heated by upwelling asthenosphere following breakoff of the Greater India slab (Hou et al., 2012; Zheng et al., 2012a,b). However, Gao et al. (2010a) have proposed a subducted sediment source for alkaline igneous rocks in southern Tibet.
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Hard collision with the Indian craton in the latest Eocene–early Oligocene (~ 34 Ma; Meng et al., 2012; van Hinsbergen et al., 2012; Tripathy-Lang et al., 2013; Fig. 2D) resulted in thrust stacking and uplift of the northern margin of the Indian plate, with attendant metamorphism (Treloar et al., 1989) and lower crustal anatexis to form adakite-like granitoids and leucogranites in southern Tibet (Bignold et al., 2006; Chen et al., 2011; Chung et al., 2003; Guo et al., 2007; Zheng et al., 2012a). At the same time (Oligo-Miocene), Tibet was also being underthrust to the north by Tarim lithosphere of the Eurasian plate, resulting in metamorphism, granite plutonism, and uplift in the Pamir (Burtman and Molnar, 1993; Robinson et al., 2007). 5.4. Neogene Despite the continued (albeit slower) convergence between India and Asia in the Miocene, the predominant Neogene structural features in Tibet and the Pamir are N–S extensional faults and grabens, reflecting gravitational collapse and E–W extension of the orogen since ~ 14 Ma (Blisniuk et al., 2001; Chen et al., 2010; Coleman and Hodges, 1995; England and Houseman, 1989; Robinson et al., 2007; Styron et al., 2013; Xu et al., 2013), or earlier (Wang et al., 2010). These extensional structures controlled the upper crustal emplacement of a suite of high-Sr/Y (adakite-like) plutons and potassic to ultrapotassic intrusive and extrusive rocks in southern Tibet (Liu et al., 2014), and leucogranites derived from partial melting of the underthrust Indian crust in the Greater Himalaya (Searle et al., 2009). Most authors identify the lower Tibetan crust or mantle as the source of Miocene magmas in Tibet (e.g., Chen et al., 2010; Wang et al., 2005), perhaps with some contribution from upwelling asthenospheric melts following rollback and breakoff of the Greater Indian slab, or melts from the underthrust Indian lithosphere (e.g., Ding et al., 2003; Guo et al., 2013; Jiang et al., 2012; R. Wang et al., 2014a; Xu et al., 2010). The debate as to the origin of these Miocene magmas is important, because they are the source of several of the largest porphyry Cu–Mo deposits in Tibet (Fig. 3), including Qulong (~16 Ma; Yang et al., 2009; Xiao et al., 2012), Jiama (~15 Ma; Ying et al., 2014), Bangpu (~15 Ma; Wang et al., 2012), and Dabu (~ 15 Ma; S. Wu et al., 2014). The high high-Sr/Y (adakite-like) character of these otherwise normal high-K calc-alkaline magmas has led some authors to propose an oceanic slab melting model for their origin (e.g., Gao et al., 2007, 2010b; Qu et al., 2004), but it seems unlikely that any oceanic lithosphere was present beneath Tibet by this time, the former Neotethyan slab having detached and sunk away into the deep mantle shortly after early Eocene soft collision (Gao et al., 2008; Hafkenscheid et al., 2006; Replumaz et al., 2010; van der Voo et al., 1999). Most recent authors concur that these magmas were generated in the Tibetan lower crust, which had been modified (underplated by mafic magmas and/or metasomatized) during Neotethyan subduction. Their adakite-like high-Sr/Y and La/Yb geochemical signatures are attributed to partial melting of these eclogitized or garnet amphibolitic lower crustal rocks (e.g., Hou et al., 2004, 2009; J.-X. Li et al., 2011). In particular, Hou et al. (2013) and S. Wu et al. (2014) have suggested that these magmas acquired their metal content by resorbing residual sulfide phases during partial melting of amphibolitic arc cumulates. Furthermore, R. Wang et al. (2014b, c) have shown that the relatively high oxidation states and water contents of these partial melts made them particularly favorable for the generation of porphyry-type magmatic–hydrothermal ore deposits upon upper crustal emplacement. R. Wang et al. (2014a) have also argued that underthrusting of the Indian continental lithosphere beneath western Tibet in the Miocene accounts for the restriction of porphyry deposits to the eastern Gangdese belt (east of ~89°E). To the west of ~89°E, Indian lithosphere was directly contacting the base of the Tibetan lithosphere by this time, displacing hot asthenospheric mantle and thereby reducing the potential for partial melting. Those magmas that are found to the west in the Miocene are mostly ultrapotassic in character, and carry isotopic
indications of derivation (at least in part) from the underthrust Indian crust (Zhao et al., 2009). Indian lithosphere now underlies the entire Gangdese belt and most of southern Tibet, while Asian lithosphere has been underthrust from the north (Kumar et al., 2006; Li et al., 2008; Nábělek et al., 2009; Zhao et al., 2011). Consequently, volcanism has largely been extinguished in this region. 6. Convergent and collisional margin tectonics, magmatism, and metallogeny The Mesozoic–Cenozoic Tethyan orogen is an ideal location to explore the transition in tectonomagmatic and metallogenic processes from the (relatively well understood) conditions encountered during oceanic lithosphere subduction to pre-, syn- and post-collisional geodynamic settings. Perhaps one of the most remarkable features of this transition is that the general characteristics of much of the magmatism and most of the resultant ore deposits do not change very much. Thus, apart from some subtle trace element characteristics (e.g., higher Sr/Y ratios in some cases), subduction-related and collisional magmas are broadly calc-alkaline to mildly alkaline (shoshonitic), with broadly similar isotopic compositions. Likewise, the nature of associated porphyry-type deposits changes little: Miocene post-collisional porphyry Cu–Mo deposits in Tibet are very similar to Mesozoic subduction-related deposits from the same region or in the Balkans (or Chile), although in some settings syn- to post-collisional deposits may be more Au-rich (e.g., Roşia Montană in the Carpathians; Kisladag and Çöpler in Turkey; Sari Gunay and Dalli in Iran). This implies that, while magmas and metals for these systems might ultimately be derived from the mantle, it is lithospheric processes, and more specifically deep crustal processes that ultimately control the compositions and metal loads of magmas reaching the upper crust during both subduction and post-subduction tectonomagmatic events. Where mantle-derived magmas are found in such settings, such as the Neogene mafic, alkaline, post-collisional volcanic rocks observed from Turkey to Tibet, these systems are generally barren (although they deserve close examination for their potential to host alkalic-type Au deposits; Richards, 1995, 2009; Jensen and Barton, 2000; A.C. Harris et al., 2013). The role of lower crustal MASH processes (melting, assimilation, storage, homogenization) in generating relatively homogeneous andesitic arc magmas by interaction between mantle-derived hydrous basalts and upper-plate crust is well understood (Annen et al., 2006; Dufek and Bergantz, 2005; Hildreth and Moorbath, 1988), and seems also to generate globally uniform porphyry Cu ± Mo ± Au and associated high-sulfidation epithermal Cu–Au deposits (Fig. 4A; Gustafson, 1979; Mitchell, 1992; Tosdal and Richards, 2001; Richards, 2003a, 2005, 2011; Candela and Piccoli, 2005; Sillitoe, 2010). For example, Mesozoic arc-related porphyry deposits from the Balkans such as Majdanpek, Elatsite, and Assarel related to subduction of the Vardar ocean (e.g., Ciobanu et al., 2002; von Quadt et al., 2005), are directly comparable to mid-Cenozoic deposits in Chile related to subduction of the Nazca plate (e.g., Sillitoe, 1989). Similarly, high-sulfidation Cu–Au deposits such as Chelopech in Bulgaria (Moritz et al., 2005) are directly comparable to deposits in the El Indio belt of Chile/Argentina (e.g., Chouinard et al., 2005). Processes that control post-subduction or post-collisional magmatism and associated ore deposits are less well understood (Fig. 4B–D). However, most researchers agree that, in the absence of a subduction-enriched asthenospheric mantle source, an origin must be sought in the lithosphere, and that lithospheric “fertilization” by precursor arc magmatism is important. Potential source regions include subduction-modified lower crust (e.g., Chiaradia et al., 2009; Richards, 2009), sub-continental lithospheric mantle (e.g., Jiang et al., 2006; Mair et al., 2011), or juvenile lower crust formed by underplated arc basalts (e.g., Hou et al., 2004, 2009; Shafiei et al., 2009). These lithospheric sources can commonly be identified from radiogenic and stable isotopic signatures (e.g., Aitcheson et al., 1995; Chiaradia et al., 2004; Jiang et al.,
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2006; Shafiei, 2010), but differ little from those of the precursor arc magmas (which themselves reflect the homogenized isotopic composition of the lower crustal MASH zone). Precursor subduction magmatism and lithospheric metasomatism is important for three reasons: (1) Aqueous metasomatism and addition of lower crustal cumulates rich in hydrous minerals such as amphibole generates lithologies that are inherently more fusible than typical anhydrous (granulitic) lower crustal assemblages; they are therefore easier to melt during later tectonic processes such as crustal thickening, or thinning with mafic magma invasion. (2) The passage of arc magmas through the crust can be expected to leave at least some metal behind in these metasomatic and cumulate zones, which can be remobilized during later partial melting. (3) Partial melts from such subductionmodified source rocks will be hydrous and oxidized (unless reduced lithologies are present in the lower crust, as perhaps in Japan; Sato, 2012; Tomkins et al., 2012), and will therefore have the potential to form magmatic–hydrothermal ore deposits upon upper crustal emplacement. It is also notable that partial melting of garnet amphibolitic sources such as those expected to be developed in the roots of continental arcs will produce magmas with high Sr/Y and La/Yb ratios (due to the presence of garnet and suppression of plagioclase crystallization; Kay et al., 1991; Haschke et al., 2002; Richards and Kerrich, 2007; Richards,
A) Subduction
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2009, 2011), as observed in many of the cases discussed here. The otherwise rather “normal” calc-alkaline to mildly alkaline compositions of these post-collisional magmas simply reflects partial melting of the residues of previous cycles of calc-alkaline arc magmatism. In contrast, low degree partial melting of metasomatized SCLM generates mafic alkaline magmas (e.g., Aydin et al., 2008; Pang et al., 2012; Pearce et al., 1990), which for the most part seem to be barren in the Tethyan orogen (although as noted above they are important targets for Au exploration in other regions; Fig. 4B, D). The role of residual sulfides in subduction-modified lithosphere is intriguing, and has been speculated upon by several authors (e.g., Chiaradia et al., 2009; Richards, 1995, 2009; Shafiei et al., 2009; Solomon, 1990; Tomkins et al., 2009). Lee et al. (2012) and Chiaradia (2014) have proposed that large amounts of Cu (up to 80% of the original arc magma flux) may be left in lower crustal arc roots, and that this residue constitutes a uniquely rich source of metals for subsequent magmatism. However, in the absence of direct evidence for such unusual concentrations of metal in exposed lower crustal arc sections, Richards (2009, 2011) has proposed that the observed trace amounts of sulfide minerals in arc cumulates from Talkeetna (Alaska) and Kohistan serve to control metal ratios (e.g., Au:Cu) in subsequent partial melts, rather than their overall metal content (Fig. 4B, D).
B) Terrane collision Collision-related porphyry Cu-Au
Arc porphyry Cu±Mo±Au Continental
Ocean
SCLM
ab n
io
at
dr
hy
de
Asthenospheric partial melting
Continental
Ocean
crust
Sl
Subducting oceanic lithosphere
MASH zone
Subducting oceanic lithosphere
Back-arc alkalic Cu-Au
Collided terrane
crust
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Asthenospheric corner flow convection Back-arc asthenospheric convection Upwelling asthenosphere Slab breakoff
C) Continental collision
Minor mafic alkaline and felsic volcanism, partly derived from underthrust or thickened lithosphere; few mineral deposits
with underthrusting
Continental crust
SML
Thickened continental crust
D) Post-collision relaxation ± delamination
Continental crust
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SML
Post-collisional alkalic Cu-Au
Thickened continental crust SCLM
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SCLM Upwelling Up asthe asthenosphere
Upwelling asthenosphere Slab breakoff
M SCLM ination delamination
Fig. 4. Schematic models of various subduction and collisional tectonic settings thought to characterize the Tethyan orogen at various times and locations, and suggested mechanisms for generating porphyry-type mineralization. (A) Normal subduction and arc magmatism generating typical calc-alkaline Cu ± Mo ± Au porphyry deposits from lower crustal melting–assimilation– storage–homogenization (MASH) zones. (B) Terrane collision (e.g., arc or microcontinent collision) with calc-alkaline Cu ± Mo ± Au porphyry deposits formed by remobilization of subduction-modified lithosphere (SML), and potential alkalic-type porphyry or epithermal Cu–Au deposits in back-arc extensional settings derived from metasomatized asthenosphere. (C) Continental collision with lithospheric underthrusting; this setting does not appear to be conducive for porphyry-type mineralization due to exclusion of hot asthenosphere from below much of the orogen; S-type magmatism from crustal melting might host lithophile element deposits. (D) Post-collisional orogenic collapse and/or delamination of subcontinental lithospheric mantle (SCLM) may bring hot asthenospheric mantle into close contact with subduction-modified lithosphere (SML), generating calc-alkaline to mildly alkaline Cu ± Mo ± Au porphyry deposits; alkalic-type porphyry or epithermal Cu–Au deposits might also potentially form in such settings, although there are few known examples in the Tethyan belt.
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7. Spatial and temporal distribution of porphyry deposits along the Tethyan orogen
which may further have reduced the likelihood of forming large porphyry deposits in this part of the belt.
The distribution of porphyry deposits along the Tethyan orogen is not uniform either in space or time (Fig. 2). Major clusters of deposits occur, in some cases of a singular age (e.g., the middle Eocene Yulong belt of Tibet, the middle Miocene deposits of southeastern central Iran), but in others with a range of ages spanning the evolution of the orogen (e.g., the Balkans, southern Tibet). There are also significant sections of the orogen that appear to be devoid of porphyry deposits, despite continuation of the tectonic and geological features that are elsewhere linked to mineralization (e.g., central Anatolia, northwestern central Iran, and northern Pakistan–Pamir–western Tibet). However, post-discovery, it is generally easier to explain why a mineral deposit has formed, rather than why a deposit has not formed in a particular area. Thus, there are numerous published explanations for the origins of the Balkan, Iranian, and Tibetan porphyry belts (reviewed above), but few studies have examined the cause of apparently barren sections of the orogen. Before seeking more complex tectonic or magmatic explanations for the apparent lack of mineralization in a particular region, the extent of cover rocks, depth of crustal exposure, and intensity of mineral exploration must first be discounted as causes. For example, it is highly likely that several more porphyry deposits could be found in central Iran by using modern geophysical exploration methods that can penetrate beneath young cover sediments, which extensively cover this region. On the other hand, intense tectonism, uplift, and erosion have removed upper crustal rocks from other areas such as northern Pakistan, while the Pamir and western Tibet are relatively unexplored. Notwithstanding the above, there may be valid tectonomagmatic reasons for the formation of discrete belts of porphyry deposits in space and time, separated by barren gaps along the orogen. Bertrand et al. (2014) have recently proposed that the irregular, clumped distribution of porphyries in the Tethyan and Andean orogens can be related to details of subduction dynamics, with folds in the downgoing slab affecting convergence rate and magma flux. They propose that arc sections characterized by a transition from rapid convergence and high magma production rates, to slower convergence and stress relaxation in the upper plate may be particularly conducive to shallow crustal plutonism and porphyry formation (e.g., Tosdal and Richards, 2001). Similarly, Cooke et al. (2005) and Rosenbaum et al. (2005) have proposed that density anomalies on the subducting plate, such as aseismic ridges, seamounts, or oceanic plateaux, can affect mantle wedge and upper plate dynamics, causing periods of compressional tectonism, crustal thickening, and fertile magmatism. In many cases, however, especially in older or collided arc systems, there is little or no record of such detailed subduction dynamics that could help guide exploration to a specific area. Rather, the occurrence of deposits is more typically used to interpret subduction history, instead of the other way around. The Tethyan orogen is also dissimilar to the Andean porphyry belt in that the majority of known deposits were formed during or after arc– continent or continent–continent collision events. Relatively few deposits (e.g., the Mesozoic Balkan and Tibetan systems) are related to simple subduction, although this is more likely to be a function of preservation than an indication that Paleo- or Neotethyan subduction was unfavorable for porphyry formation. Nevertheless, there is an intriguing possibility that the almost complete absence of porphyry deposits in the western Tethys (Spain and the Alps) reflects Early Cretaceous ocean anoxia (Westermann et al., 2013). Subduction of reduced seafloor sediments and unoxidized oceanic crust would not be optimal for forming oxidized arc magmas and associated porphyry Cu ± Mo ± Au deposits (Richards and Mumin, 2013). In addition, because closure of the Tethyan ocean basins occurred by anticlockwise pivotal rotation of Africa towards Eurasia about an axis close to Spain (Fig. 1; Schettino and Turco, 2011), there has been relatively little subduction (in terms of volume of subducted oceanic lithosphere) in the western Tethys,
7.1. Post-collisional porphyry deposits Until relatively recently, collisional orogens were not thought to be prospective for porphyry-type mineralization, based on the assumption that active subduction was required to generate the requisite fertile magmas. Indeed, the occurrence of porphyry deposits (and associated calc-alkaline magmatism) was commonly used to constrain tectonic histories, on the basis that the deposits must be subduction-related, and therefore must precede any collisional events. The discovery of numerous normal-looking porphyry Cu deposits in Tibet over the last decade (e.g., Hou et al., 2004), and their accurate dating, as well as detailed reconstruction of the tectonic history of the Indian–Asian collision, have demonstrated convincingly that these deposits were post-collisional, and likely formed 30–40 m.y. after the termination of oceanic lithosphere subduction (Hou et al., 2009, 2011). Similarly, more detailed geochronological, geochemical, and structural studies in southeastern Europe, Turkey, and Iran have revealed that many deposits previously assumed to be subduction-related are in fact syn- or postcollisional (e.g., Shafiei et al., 2009), or were formed during complex tectonic readjustments in advance of collisional events (e.g., Imer et al., 2013; Moritz et al., 2010). It is now clear that in fact the majority of Cenozoic porphyry deposits in the Tethyan orogen are back-arc, postsubduction, or collisional in origin. Thus: in the Balkans and Greece, Cenozoic porphyry and epithermal deposits have formed in a complex arc and micro-continent collisional environment relating to closure of the Vardar ocean basin; in Turkey, similar deposits have formed in response to closure of several small Neotethyan basins, with tectonic settings ranging from back-arc to syn-collisional, and post-collisional extension; and in Iran and Tibet, middle Miocene porphyry deposits are thought to have formed in response to post-collisional lithospheric melting. The distribution of these collision-related porphyry and epithermal deposits is more obviously related to particular (collided) sections of the orogen. Thus, they are quite evenly scattered, albeit with some local clustering, along the Balkan–Turkish–Iranian segments of the Afro-Arabian–Eurasian collisional orogen. The oldest (Paleogene) deposits occur in the Balkans and around the Arabian indentor in Turkey, whereas younger (Miocene) deposits occur to the southeast in Iran due to the diachronous nature of the Arabian collision. Mineralization also continued through the Miocene in the Balkans, Greece, and western Turkey in response to post-Vardar closure and Aegean extension. In southern Tibet, the Miocene collision-related porphyry Cu–Mo ± Au belt appears to be of limited longitudinal extent (~500 km), terminating westward at ~89°E. The apparent absence of Cenozoic porphyry deposits from here to western Pakistan is intriguing, but as noted above, may at least in part reflect preservational issues, as well as lack of detailed exploration. Nevertheless, R. Wang et al. (2014a) have suggested that underthrusting of the cool Indian plate beneath western Tibet in the Miocene may have limited partial melting of the overthrust Tibetan lithosphere, and hence the formation of fertile magmas. To the east, asthenospheric upwelling following slab breakoff may have generated fertile partial melts from previously subduction modified (hydrated, oxidized) Tibetan lithosphere (R. Wang et al., 2014b,c). These considerations indicate that, while collisional orogens can now be considered favorable targets for porphyry and epithermal deposit exploration, local scale (relative to an entire orogen; i.e., 500–1000 km-scale) variations in collisional history and precursor geology can significantly affect the potential for ore formation. In particular, advanced continental collision and lithospheric underthrusting does indeed appear to result in less favorable conditions for fertile magma generation (as previously thought; Fig. 4C). However, arc collisions and tectonic adjustments in advance of or following continental collision (especially where these involve transpressional or
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transtensional tectonics) can trigger remobilization of previously subduction-modified lithosphere, and the formation of porphyry and epithermal deposits that closely resemble “normal” subduction-related deposits (Fig. 4B, D).
8. Conclusions The Tethyan orogen stretches for over 12,000 km from the Alps to the Himalayas, and beyond, and hosts numerous small to world-class porphyry Cu ± Mo ± Au and related epithermal Au ± Cu deposits. Until relatively recently, it was assumed that most of these deposits were formed by the subduction of oceanic lithosphere and the generation of hydrous, oxidized arc magmas. While this appears to be true for the relatively small number of Mesozoic deposits preserved in the belt, which formed by subduction of Neotethyan oceanic lithosphere, this does not now appear to be the case for the majority of Cenozoic deposits. Improved geochronological constraints, both on the ore deposits themselves and on regional tectonic history, have shown over the past decade or so that many of these deposits in fact postdate oceanic subduction, and many have formed during or after collisional orogenic events. The mid-Miocene porphyry Cu–Mo deposits in southern Tibet are the clearest examples of post-collisional systems, post-dating subduction and slab breakoff by ≥30 m.y. Despite the complex and disputed history of the Indian–Asian collision, this time gap is too large to allow for a direct subduction involvement in magmatism and ore formation in any reasonable tectonic reconstructions. These magmatic–hydrothermal systems instead appear to have formed by remobilization of deep lithosphere modified by prior subduction processes (e.g., SCLM metasomatism and addition of cumulates and plutons to the lower arc crust). Processes that can cause this remobilization by partial melting include crustal thickening, crustal thinning, and SCLM delamination, commonly including heating by upwelling asthenospheric melts. Elsewhere along the Tethyan orogen the situation is not so clear, being complicated by the fact that closure of the Tethyan ocean was not a single event. The earliest closure events that eliminated the old Paleotethys lithosphere were concluded in the Paleozoic, and few deposits related to this subduction and collision event are preserved (Pulang and Yangla in the Indosinian belt are examples). After a pulse of Cretaceous arc magmatism and associated porphyry deposits (e.g., in the Balkans, Caucasus, and central Tibet), the Cenozoic was characterized by back-arc and syn- or post-collisional magmatism, generated by complex interactions between the former Cimmerian continental fragments and the Eurasian plate margin. The opening of numerous small back-arc basins and closure of strands of the Neotethyan ocean over small geographic distances and short timescales led to a complex geological record in which it is difficult to ascribe particular igneous sequences to a specific subduction or collisional events. Nevertheless, carefully constrained tectonic reconstructions and improved geochronology have led to the conclusion that most of the porphyry and epithermal deposits formed during this period in the western and central parts of the orogen were not formed directly from subduction processes, but reflect tectonomagmatic adjustments to the onset of collision. Some deposits appear to be truly postcollisional, as in Tibet, and may reflect lithospheric responses such as SCLM delamination or extensional orogenic collapse (as in Anatolia).
Acknowledgments I thank Zengqian Hou for inviting me to write this review paper, and Ali Imer, Amir Razavi, Rui Wang, Jing-jing Zhu, and Jinxiang Li for collaborations on various projects throughout the Tethyan belt. Yongjun Lu and an anonymous reviewer are thanked for helpful comments on the manuscript. Christopher Scotese is thanked for permission to reproduce his plate reconstruction maps in Fig. 1. This work was supported by a
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