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Earth-Science Reviews 133 (2014) 62–93

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Earth-Science Reviews

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Zircon dating of Neoproterozoic and Cambrian ophiolites in West Mongolia and implications for the timing of orogenic processes in the central part of the Central Asian Orogenic Belt

Ping Jian a,⁎, Alfred Kröner b,Bor-mingJahnc,BrianF.Windleyd, Yuruo Shi a, Wei Zhang a, Fuqin Zhang e, Laicheng Miao e,DondovTomurhuuf,DunyiLiua a SHRIMP Center, Institute of Geology, Chinese Academy of Geological Sciences, Baiwanzhuang Road 26, Beijing 100037, China b Institut für Geowissenschaften, Universität Mainz, D-55099 Mainz, Germany c Department of Geosciences, National Taiwan University, P.O. Box 13-318, Taipei 106, Taiwan d Department of Geology, University of Leicester, Leicester LE1 7RH, UK e Institute of Geology and Geophysics, Chinese Academy of Sciences, Beijing 100029, China f Institute of Geology and Mineral Resources, Mongolian Academy of Sciences, Ulaanbaatar 210351, Mongolia article info abstract

Article history: We present new isotopic and trace element data to review the geochronological/geochemical/geological Received 5 September 2012 evolution of the central part of the Central Asian Orogenic Belt (CAOB), and find a fundamental geological Accepted 27 February 2014 problem in West Mongolia, which has traditionally been subdivided into northwestern early Paleozoic Available online 6 March 2014 (formerly Caledonian) and southerly late Paleozoic (formerly Hercynian) belts by the Main Mongolian Lineament (MML). We resolve this problem with SHRIMP zircon dating of ophiolites and re-evaluation Keywords: – Ophiolite of much published literature. In Northwest Mongolia the Dariv Khantaishir ophiolite marks the boundary – Subduction initiation between the Lake arc in the west and the Dzabkhan Baydrag microcontinent in the east. Zircons from a 206 238 Arc–microcontinent collision microgabbro and four plagiogranites yielded weighted mean Pb/ U ages of 568 ± 5 Ma, 567 ± 4 Ma, Subduction–polarity reversal 560 ± 8 Ma (Dariv), 573 ± 8 Ma and 566 ± 7 Ma (Khantaishir) that we interpret as reflecting the time of Central Asian Orogenic Belt ophiolite formation (ca. 573–560 Ma). Metamorphic zircons from an amphibolite on a thrust boundary between Zircon dating the Khantaishir ophiolite and the Dzabkhan–Baydrag microcontinent formed at 514 ± 8 Ma, which we interpret as the time of overthrusting. In South Mongolia the Gobi Altai ophiolite and the Trans-Altai Gurvan Sayhan– Zoolen forearc with an ophiolite basement were investigated. Zircons of a layered gabbro (lower ophiolite crust) and a leucogabbro (mid-upper crust) of the Gobi Altai ophiolite yielded crystallization ages of 523 ± 5 Ma and 518 ± 6 Ma. The age data constrain the formation time of ophiolite within ca. 523–518 Ma. Zircons from four samples of the Gurvan Sayhan–Zoolen forearc, with similar hybrid adakite–boninite affinities, yielded 519 ± 4 Ma for an anorthosite, ≥ 512±4Maforahornblendite and 520 ± 5 and 511 ± 5 Ma for two diorites. The ophiolite basement has an upper age limit of 494 ± 6 Ma, determined by dating a tonalite dike cutting the Zoolen ophiolite. Integrating available zircon ages as well as geochemical and geological data, we re-subdivide West Mongolia into: a latest Neoproterozoic-early Cambrian, arc–microcontinent collision zone north of the MML; a Cambrian Gobi Altai ophiolite–microcontinent collision zone and a Cambrian Trans–Altai forearc complex south of the MML. The central CAOB evolved in five phases: subduction initiation and arc formation (ca. 573 to N ca. 540 Ma); arc–microcontinent collision (ca. 535–524 Ma); a continuum of slab delamination, overthrusting, crustal thickening and surface uplift (ca. 519–482 Ma) in Northwest Mongolia; initiation of new subduction zones in South Mongolia (ca. 523–511 Ma); and continuing orogeny with local surface uplift. Overall, the current, documented timing of orogenic development in the central CAOB is largely consistent with a W/SW-Pacific style of evolution in terms of subduction initiation, short timescales of individual orogenies, and episodic subduction–collision during a continuing migration of subduction zones. © 2014 Elsevier B.V. All rights reserved.

⁎ Corresponding author. Tel.: +86 10 68999765; fax: +86 10 68311545. E-mail address: [email protected] (P. Jian).

P. Jian et al. / Earth-Science Reviews 133 (2014) 62–93 63

Contents

1. Introduction...... 63 2. Regionalgeology,publishedzircondataforophiolitesandafundamentalproblemintectonicsubdivision...... 64 2.1. Regionalgeology...... 64 2.1.1. NorthwestMongolia...... 64 2.1.2. SouthMongolia...... 67 2.2. Publishedzircondatafortheophiolites...... 69 2.3. Afundamentalproblemintectonicsubdivision...... 69 3. Approachtoconstraintheformationofophiolites...... 69 3.1. Appropriatenessofgabbroicandleucocraticrockstozirconchronology...... 69 3.1.1. Gabbroicrocks...... 69 3.1.2. Leucocraticrocks(oceanicplagiogranites)...... 70 3.2. Methodology...... 70 4. Fieldinvestigationandsamples...... 70 4.1. DarivandKhantaishir(NorthwestMongolia)...... 70 4.1.1. Dariv...... 70 4.1.2. Khantaishir...... 70 4.2. GobiAltai(SouthMongolia)...... 70 4.3. GurvanSayhanandZoolen(Trans-Altaizone,SouthMongolia)...... 73 4.3.1. GurvanSayhan...... 73 4.3.2. Zoolen...... 73 5. Zirconages...... 73 5.1. DarivandKhantaishir(NorthwestMongolia)...... 73 5.1.1. Dariv...... 73 5.1.2. Khantaishir...... 73 5.1.3. Summary...... 76 5.2. GobiAltai(SouthMongolia)...... 77 5.2.1. Layeredgabbrofromanorthernthrustslice...... 77 5.2.2. Isotropicleucogabbrofromasouthernvolcano-sedimentarymélange...... 77 5.2.3. Summary...... 78 5.3. GurvanSayhanandZoolen(Trans-Altai,SouthMongolia)...... 78 5.3.1. Anorthosite(patch)andhornblendite(tectonicblock)inavolcano-sedimentarymélange(GurvanSayhan)...... 78 5.3.2. Diorites(tectonicblocks)inmetavolcano-sedimentarymélanges(Zoolen)...... 79 5.3.3. Tonalitedikes(Zoolen)...... 80 5.3.4. Summary...... 81 6. Geochemicaldata...... 81 6.1. DarivandKhantaishir(NorthwestMongolia)...... 81 6.1.1. Dariv...... 81 6.1.2. Dioriteandleucogranitedikes...... 83 6.1.3. Khantaishir...... 83 6.1.4. Nd–Srisotopiccompositionsofselectedsamples...... 83 6.2. GobiAltai(SouthMongolia)...... 85 6.3. GurvanSayhanandZoolen(Trans-Altai,SouthMongolia)...... 85 6.3.1. Metaperidotites(serpentinites)...... 85 6.3.2. Hornblendite–anorthosite(atectonicblock)...... 85 6.3.3. Diorites(tectonicblocks)andadiabasedike...... 85 6.3.4. Atonalitedikecuttingophiolite...... 88 6.3.5. Nd–Srisotopiccompositions...... 88 7. Discussion:timingoforogenyinthecentralCAOB...... 88 7.1. Phase 1 (ca. 573 to N ca.540Ma):subductioninitiationandarcformationinNorthwestMongolia...... 88 7.2. Phase 2 (ca. 535–524Ma):likelysyn-collisiongranitoidmagmatismintheLakearc...... 88 7.3. Phase 3 (ca. 519–482Ma):acontinuumofslabdelamination,overthrusting,crustalthickeningandsurfaceuplift...... 89 7.4. Phase 4 (ca. 523–511Ma):initiationofnewsubductionzone(s)inSouthMongolia...... 89 7.5. Phase5:continuingorogenywithlocaluplift...... 89 8. Conclusions...... 90 Acknowledgments...... 90 References...... 90

1. Introduction et al., 1993; Davies and von Blanckenburg, 1995; Turner et al., 1999; Rey et al., 2001; Jadamec et al., 2007). The plate tectonic paradigm or Wilson cycle (Wilson, 1966) begins Accretionary orogens form at intra-oceanic and/or (rifted) with continental rifting, and proceeds via ocean-ﬂoor spreading and continental-margin convergent plate boundaries and often are the subduction, to ocean closure and terminates with continental collision. locus of major juvenile crustal growth (e.g., Hamilton, 1988; Windley, This cycle explains the geological evolution of classic collisional orogens 1992; Jahn et al., 2000; Shervais, 2001; Jahn et al., 2004; Shervais et al., such as the Alpine–Himalayan chain, which resulted from the closure of 2004; for overview see also Cawood et al., 2009). A conceptual difference Neo-Tethys. The termination of orogeny, i.e. cessation of contractional between an accretionary and a collisional orogen is that the former gen- continental deformation, is usually signiﬁed by surface uplift due to ex- erally lacks the dominant role of a colliding continent (e.g., Windley, tensional collapse of over-thickened crust–mantle (Bird, 1979; Dewey 1992; Cawood and Buchan, 2007) whereas the latter requires Author's personal copy

64 P. Jian et al. / Earth-Science Reviews 133 (2014) 62–93 continental collision during its development. The best-known Cenozoic either the CAOB formed through continuous, long-lived subduction– accretionary orogens are the intra-oceanic West (W)-Pacific(e.g.,Stern accretion within a single major ocean (e.g., Sengör et al., 1993; and Bloomer, 1992; Stern, 2002, 2004), the (rifted Australian) Yakubchuk, 2004), or by punctuated, semi-continuous, sequential continental-margin Southwest (SW)-Pacific east and north of Australia subduction–accretion–collision processes in an oceanic archipelago (e.g., Hall, 2002, 2009; Schellart et al., 2006; Whattam et al., 2008; (e.g., Kozakov et al., 1999; Buchan et al., 2002; Kozakov et al., 2002, Whattam, 2009), and the Caribbean (Pindell and Kennan, 2009). The 2003; Jian et al., 2008, 2010a,b; Lehmann et al., 2010). first two have served as modern analogs to ancient accretionary orogens, The primary objective of this study is to analyze the thermo-tectonic such as the Tasmanides of eastern Australia (e.g., Collins, 2002; Glen development of the central CAOB in West Mongolia and adjacent areas and Meffre, 2009; Glen et al., 2009) and the Central Asian Orogenic in Chinese Inner Mongolia and Xinjiang (Fig. 1A,B) by constraining the Belt (CAOB) (Jahn et al., 2000, 2004; Kröner et al., 2007; Windley timescale of orogeny. We first review the geology of a large region et al., 2007; Kröner et al., 2010) or Altaids (e.g., Sengör et al., 1993; that is bounded to the north by the Neoproterozoic (ca. 655–636 Ma; Yakubchuk et al., 2001; Yakubchuk, 2002, 2004; Wilhem et al., 2012). Jian et al., 2010a), MORB (Mid-Ocean-Ridge-Basalt)-type Bayankhongor Using the Cenozoic W- and SW-Pacific as an analog to account for ancient ophiolite (Fig. 1A), to the northwest by the Cambrian–Ordovician accretionary orogens is insightful, in particular, in the following aspects: (ca. 503–481 Ma; Jian et al., 2003, 2005; Xiao et al., 2006)Zhaheba ophiolite (Fig. 1A) in north Xinjiang, China, and to the southeast by (1) Subduction initiation—the W-Pacific Izu–Bonin–Mariana (IBM) the Permian (ca. 288–284 Ma) SSZ-type Solonker ophiolite (Jian and SW-PacificTonga–Kermadec (TK) intra-oceanic convergent et al., 2010b) along the Mongolia/China border (Fig. 1A). We then margins are type localities of supra-subduction zone (SSZ) focus on the ophiolites and mélanges north and south of the Main ophiolites that marked mid-Eocene subduction initiation, Mongolian Lineament (MML) (Fig. 1A), which is a crustal-scale fault which was likely responsible for a global reorganization of plate zone believed to represent a regional structural boundary (Tomurtogoo, boundaries (Stern and Bloomer, 1992; Stern, 2004; Stern et al., 1997, 2002; Xiao et al., 2004; Windley et al., 2007). 2012) or a global-scale plate interaction (Flower et al., 2001). (2) Short timescales of individual orogenies (e.g., Dewey, 2005)—most 2. Regional geology, published zircon data for ophiolites and a SSZ-type ophiolites (forearc basements) and backarc basins fundamental problem in tectonic subdivision have relatively short life-cycles (10–30 Ma; mostly b20 Ma) (e.g., Shervais, 2001; Collins, 2002; Schellart et al., 2006), 2.1. Regional geology and the IBM forearc (e.g., Stern, 2002, 2004) and the Parece– Vela backarc basin (e.g., Sdrolias et al., 2004) in the W-Pacific We refer to the region north of the MML as Northwest Mongolia are excellent examples. and in the south to South Mongolia that extends to the Chinese Altai (3) Episodic subduction–collision during continuing orogeny—the SW- (Fig. 1A). Following Kröner et al. (2010) and Lehmann et al. (2010), Pacific contains a wide range of tectonic elements that include we use non-genetic terms such as (tectonic) “zone”, “block” and continental ribbons, SSZ-type ophiolites, island arcs, oceanic “belt”,todefine the principal petrotectonic units. islands/plateaus and back-arc basins, and it preserves evidence of Cenozoic to Recent collisions between an arc (usually with a 2.1.1. Northwest Mongolia SSZ-type ophiolite basement) and a microcontinent or an ocean- Northwest Mongolia comprises three fault-bound units from east to ic island/plateau (Hall, 2002; Whattam et al., 2008; Hall, 2009; west, the Dzabkhan–Baydrag block (microcontinent), the Lake zone Whattam, 2009). Such local collisions are best manifested by (arc) and the Hovd zone. As shown in Fig. 1A, a prominent feature of obduction of SSZ-type ophiolites onto continental ribbons origi- this region is the presence of a possible arc–microcontinent collision nally rifted from Australia and often result in a subduction polar- zone and an ophiolite-decorated suture. ity flip and further orogenic development (e.g., Stern, 2004; Whattam et al., 2008; Whattam, 2009). 2.1.1.1. Dzabkhan–Baydrag block (microcontinent). This is a micro- (4) Seaward migration of subduction zones—the W-Pacificand continent (Fig. 1A) with an Archean-Mesoproterozoic crystalline base- SW-Pacific east of Australia are modern retreating orogens ment and a Mesoproterozoic-Neoproterozoic sedimentary cover (Kotov (e.g., Collins, 2002; Cawood et al., 2009) that are characterized et al., 1995; Kozakov et al., 2001, 2007; Demoux et al., 2009a,b). The by episodic slab rollback and seaward migration of subduction microcontinent underwent two episodes of crustal reworking and zones (e.g., Lonergan and White, 1997; Schellart et al., 2006; associated plutonism to form granitic gneiss/migmatite complexes and Schellart and Rawlinson, 2010). diorite–granite intrusions, the earlier in the mid-Neoproterozoic (ca. The Central Asian Orogenic Belt (CAOB) occupies a broad area 755–856 Ma; SHRIMP, ICP-MS and TIMS zircon ages; Zhao et al., 2006; between the Siberian, Baltic, North China and Tarim cratons (Fig. 1B) Yarmolyuk et al., 2008a; Levashova et al., 2010; Kozakov et al., 2012) and, like the Cenozoic SW-Pacific, it contains a wide range of tectonic el- and the later in the Early Paleozoic (ca. 515–482 Ma; Dijkstra et al., ements that include ophiolites, arc–back arc systems, oceanic islands/ 2006; Kröner et al., unpublished SHRIMP zircon, for analytical data see plateaus, subduction–accretion complexes and microcontinents Table S2; Fig. 2A,B). Note that the later intrusions were broadly synchro- (e.g., Sengör et al., 1993; Jahn et al., 2000, 2004; Kozakov et al., 2001; nous with a local phase of granulite-facies metamorphism (ca. 510 Ma; Sengör and Natal'in, 2004; Kröner et al., 2007; Windley et al., 2007). Fig. 2A; Kozakov et al., 2002) in a gneiss-dome (e.g., Fedorovskii and The similarity in tectonic elements of the CAOB and SW-Pacific, in a Khain, 1995). This is highly relevant to the present study, which we will broad sense, may imply similar subduction–accretion–collision pro- discuss further below. cesses. However, unlike the brief Late Mesozoic-Cenozoic development in the SW-Pacific(e.g.,Hall, 2002, 2009), the CAOB was extraordinarily 2.1.1.2. Ophiolites. Along the contact between the Lake zone to the west long-lived with a total duration of presumably ca. 800 Ma (e.g., Khain and the Dzabkhan–Baydrag block to the east is a belt of ophiolite nappes et al., 2002; Kröner et al., 2007). Such a long evolution, if it is quasi- (Fig. 1A). Detailed mapping at Dariv (Fig. 2A; Khain et al., 2003; Dijkstra continuous as many researchers believe, is unusual and is in conflict et al., 2006) and Khantaishir (Fig. 2B; Zonenshain and Kuzmin, 1978) with the corollary that an individual orogeny normally has a short dura- demonstrated that the ophiolites generally dip NW and are thrust tion (Dewey, 2005), as documented in the Cenozoic W- and SW-Pacific onto the Dzabkhan–Baydrag block. orogens (e.g., Stern, 2002, 2004; Sdrolias et al., 2004)aswellasinmany 2.1.1.2.1. Dariv. The Dariv ophiolite in the northwestern part of the ancient orogens (Dewey, 2005). A current debate concerning the timing Dariv Range has a near-complete ophiolite stratigraphy (Fig. 2A), of orogeny is thus focused on two contrasting evolutionary models: and its petrology (e.g., Khain et al., 2003), geochronology (e.g., Author's personal copy

P. Jian et al. / Earth-Science Reviews 133 (2014) 62–93 65

Fig. 1. (A) Geological sketch map of western Mongolia and adjacent areas in China (compiled after GSIGMR, 1998; Badarch et al., 2002; Xiao et al., 2004; Windley et al., 2007; Jian et al., 2008; Kröner et al., 2010). Note that the Main Mongolian Lineament (MML) can be deﬁned differently, and here the lineament is after Xiao et al. (2004) (for the Chinese part) and Windley et al. (2007) (for the Mongolian part). SHRIMP zircon ages (literature data and our new data) of ophiolites and relevant rocks are marked. For a Google map corresponding to Fig. 1A see appendix Fig. 1A. (B) Map showing the tectonic framework of Asia (Jian et al., 2010b) with principal cratons (Siberian, Baltica, Tarim, North China, Yangtze and Indian), orogens (CAOB, China Central and Paleo-Tethys/Tethys), the Solonker (Permian) suture (red line) and the position of Fig. 1A. Also shown are the Siberian (SCFB) and the Emeishan (ECFB) ﬂood basalt provinces.

Kozakov et al., 2002; Khain et al., 2003) and geochemistry (Dijkstra peridotites (harzburgite–dunite) (Khain et al., 2003; Dijkstra et al., et al., 2006) have been studied. It has the following downward 2006). Based on the compositions of the lavas and dikes, some of stratigraphy: intermediate to felsic lavas, sheeted dikes, maﬁcand which are boninitic, Dijkstra et al. (2006) suggested that the Dariv ultramaﬁc cumulates (gabbro–norite, websterite, orthopyroxenite; ophiolite represents a proto-arc formed above a nascent subduction i.e. the gabbro–pyroxenite in Fig. 2A), and serpentinized mantle zone. Author's personal copy

66 P. Jian et al. / Earth-Science Reviews 133 (2014) 62–93

Fig. 2. Geological sketch maps of the Dariv (A) (modiﬁed after Khain et al., 2003; Dijkstra et al., 2006) and Khantaishir (B) (modiﬁed after Zonenshain and Kuzmin, 1978) areas. Relevant zircon data are compiled from the literature, and sample localities are indicated. For Google maps corresponding to Figs. 1A and B see appendix Figs. 1AandB.

2.1.1.2.2. Khantaishir. The Khantaishir ophiolite near Altai City has Matsumoto and Tomurtogoo, 2003). Geochemical data of the the following interpreted upwards stratigraphy: harzburgite–dunite, sheeted dikes (predominately boninitic), lavas (tholeiitic and calc- gabbro–pyroxenite, layered and isotropic gabbro, sheeted dikes, pil- alkaline) and gabbro–pyroxenites (mostly tholeiitic) indicate a SSZ origin low lava, and chert (Fig. 2B); most rocks occur in a serpentinite mé- (Zonenshain and Kuzmin, 1978; Matsumoto and Tomurtogoo, 2003;for lange (Zonenshain and Kuzmin, 1978; Zonenshain et al., 1985; overview see also Dergunov, 2001). Author's personal copy

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2.1.1.3. Lake zone (predominantly a volcanic arc). TheLakezone(Fig. 1A) eclogite lenses and numerous ophiolitic fragments. From structural mostly comprises volcano-sedimentary rocks, subordinate granitoids studies Lehmann et al. (2010) and Štípská et al. (2010) reported that (Badarch et al., 2002; Rudnev et al., 2009; Kovach et al., 2011; the Neoproterozoic continental basement is partly overlain by the sub- Yarmolyuk et al., 2011; Rudnev et al., 2012) and a minor peridotite– duction–accretion complex, which itself is tectonically overlain by a pyroxenite–anorthosite–gabbronorite (intrusive) suite (511 ± 7 Ma; fragmented ophiolite. Rudnev et al., 2009; Yarmolyuk et al., 2011). Kovach et al. (2011) 2.1.2.1.1. Neoproterozoic continental basement and an early Paleozoic and Yarmolyuk et al. (2011) distinguished two geochemical suites granitic pluton. The continental basement is the structurally lowest of volcanic rocks: a relatively low-Ti, calc-alkaline basalt–andesite– unit in the Gobi Altai zone and is represented by coarse-grained granit- dacite–rhyolite (BADR), and a relatively high-Ti, alkaline trachybasalt– oids mostly augen-gneiss, hornblende gneiss, amphibolite, marble and trachybasaltic andesite–trachyandesite. Volcanic rocks of the BADR metapelite (Kröner et al., 2010). SHRIMP zircon dating of the gneisses suite are interbedded with siliceous (chert and siltstone) and clastic (Demoux et al., 2009b; Kröner et al., 2010) suggests a formation age (mostly sandstone) metasediments and are unconformably overlain by range of ca. 950–1000 Ma. The basement rocks shown in Fig. 3 are conglomerates (e.g., in Fig. 2A). The volcanic and sedimentary rocks intruded by small bodies of variably foliated diorite (542 ± 4 Ma; ICP- are similar to those of a typical island arc (e.g. Bailey, 1981; MS zircon; Hanžl and Aichler, 2007), granite (518 ± 5 Ma, ICP-MS zir- Macdonald et al., 2000). The volcanic rocks show strong negative con; Hrdličková et al., 2008) and a leucogranite dike (510 ± 6 Ma,

Nb–Ta–Ti anomalies and εNd (t = 570 Ma) values of 9.9–7.3 (Kovach SHRIMP zircon; Kröner et al., 2010). In an eastern continuation et al., 2011). Interestingly, the εNd (t = 570 Ma) values (7.4–5.6) of the (i.e. east of Bayanig, Fig. 3) of the continental basement there are siliceous-clastic metasediments are also highly positive (Kovalenko gneissic granodiorites–granites of slightly younger emplacement age et al., 2004; Kovach et al., 2011). The geochemical data, and in particular (502 ± 6–498 ± 3 Ma; SHRIMP zircon; Demoux et al., 2009b). the Nd isotopic compositions, led Kovach et al. (2011) to conclude that About 15 km east of Chandman (Fig. 3) an early Paleozoic granitic the BADR suite forms part of an intra-oceanic magmatic arc in the Lake pluton (zircon, 511 ± 5 Ma; monazite, 506 ± 5 Ma; ICP-MS) is insignif- zone. icantly deformed (Hrdličková et al., 2010). The granitic rocks are high-K The age of the Lake arc has not yet been precisely defined. However, calc-alkaline, subaluminous to slightly peraluminous, and have an initial 87 86 an andesite has an amphibole Ar–Ar age of 546 ± 3 Ma (Yarmolyuk Sr/ Sr ratio of 0.7064 and a range of negative εNd (t) values (−1.5 to et al., 2011), which in principle reflects cooling of magma through the −0.2; Hrdličková et al., 2010). blocking temperature of hornblende (ca. 500 °C; Harrison and Geraid, 2.1.2.1.2. Subduction–accretion complex with eclogite lenses. This small 1986). Moreover, according to Kröner et al. (2001), a quartz porphyry subduction–accretion complex (known as the Tsakhir Uul Formation) clast from a conglomerate at Dariv (Fig. 2A) has a zircon evaporation in the Gobi Altai zone (Fig. 3) consists essentially of a micaschist 207Pb/206Pb age of 540 ± 1 Ma. These dates suggest the intra-oceanic (matrix) mélange, which encloses lenses/blocks of eclogite (garnet + magmatic arc has an upper age limit (ca. 540–546 Ma) near the omphacite + phengite), metapelite and carbonate (Lehmann et al., Ediacaran/Cambrian boundary (542 ± 1 Ma). 2010; Štípská et al., 2010). The eclogite protoliths are geochemically Geochemical and zircon age data for several granitoids in the Lake transitional between E-MORB (enriched Mid-Ocean-Ridge Basalt) and zone (Rudnev et al., 2009; Kovach et al., 2011; Yarmolyuk et al., OIB (Ocean Island Basalt), and they were metamorphosed at peak 2011; Rudnev et al., 2012)define four discrete plutonic episodes: pressure–temperature (PT) conditions of 20–22.5 kbar and 590–610 °C (1) predominantly high-Al diorite–tonalite–trondhjemite (the oldest (Štípská et al., 2010). A phengite in eclogite yielded an Ar–Ar age of is 551 ± 13 Ma; mostly, 535 ± 6–524 ± 10 Ma; εNd(t) =9.0–7.4); 543 ± 4 Ma (1 sigma error), and white micas from two mylonitic (2) mostly low-Al diorite–tonalite–trondhjemite (519 ± 8–494 ± orthogneissess (Lehmann et al., 2010)andamicaschist(Štípská et al.,

10 Ma; εNd(t) =7.9–6.5); (3) alkaline granite–monzodiorite (511 ± 2010)haveAr–Ar ages of 573 ± 15, 540 ± 15 Ma and 537 ± 3 Ma 2Ma;εNd(t) =7.2–6.0); and (4) undeformed diorite–granodiorite– respectively. These Ar–Ar ages may suggest exhumation of the sub- granite–leucogranite (465 ± 11–456 ± 4 Ma; εNd(t) =6.9–3.1). Nota- duction–accretion complex in the latest Neoproterozoic-earliest bly, all these granitoids have variable but high positive εNd(t) values. Cambrian (ca. 573–537 Ma). However, phengite and white mica in high-pressure metamorphic rocks may contain excess Ar, and such Ar– 2.1.1.4. Hovd zone. The Hovd zone (Fig. 1A) is characterized by extensive Ar ages may therefore be geochronologically doubtful (Li et al., 1993, Ordovician–Silurian clastic-volcanogenic sequences and reef limestones 1994). (Markova, 1975; Dergunov, 1989; Badarch et al., 2002). Greenschist- 2.1.2.1.3. Ophiolite. The Gobi Altai ophiolite contains, from bottom facies volcano-sedimentary rocks of unknown age, confined to a narrow to top, serpentinized peridotite, gabbro, diabase dikes, tonalite– tectonic wedge on the eastern margin of the Lake zone (Geological Map trondhjemite, and tholeiitic to calc-alkaline mafictointermediate of Western Mongolia, 1990; Geological Map of Mongolia, 1999), are in- volcanic rocks (Hanžl and Aichler, 2007). However, the original stratig- truded by several layered picrite–peridotite–gabbro massifs of likely raphy is tectonically dismembered. As shown in Fig. 3, several ophiolitic Cambrian age (e.g., 512 ± 6 Ma, 39Ar–40Ar biotite age; Izokh et al., fragments are in tectonic contact with the Neoproterozoic basement 2010, 2011). rocks, and others occur in a sheared metavolcano-sedimentary matrix of possible Cambrian–Ordovician age (Fig. 3). In order to determine 2.1.2. South Mongolia the depositional age of the sheared matrix, Kröner et al. (2010) reported South Mongolia includes the Gobi Altai, Trans-Altai and South Gobi detrital zircon ages of ca. 467–538 Ma and ca. 1909–2672 Ma from three zones (Fig. 1A), each of which contains distinctive rock assemblages sandstone samples (SHRIMP data; Fig. 3). Also, a metarhyolite has an and tectonic features. Below we describe the ophiolitic fragments in eruption age of 520 ± 5 Ma (ICP-MS zircon; Kröner et al. unpublished; detail in the Gobi Altai (Fig. 3) and Trans-Altai (Fig. 4) zones and briefly for analytical data see Table S2; Fig. 3) and a metadacite has an eruption the geology of the South Gobi zone. The three zones likely belong to a time of 443 ± 5 Ma (SHRIMP zircon; Kröner et al. unpublished; for an- continental block that consolidated at ca. 498–420 Ma; (Jian et al., alytical data see Table S2; Fig. 3). 2008) on the northern side of the Permian ophiolite-decorated Solonker suture (Fig. 1A,B; Xiao et al., 2003; Jian et al., 2010b). 2.1.2.2. Trans-Altai zone: the Gurvan Sayhan and Zoolen ophiolitic mélanges. The Trans-Altai zone (Kröner et al., 2010; Lehmann et al., 2.1.2.1. Gobi Altai zone. The narrow Gobi Altai zone (Fig. 3)onthe 2010) contains at Gurvan Sayhan and Zoolen ophiolite mélanges (Fig. 4) southern side of the MML (Fig. 1A) includes a Neoproterozoic continen- first described by Zonenshain (1973). The ophiolites form numerous frag- tal basement that we call the Gobi Altai microcontinent, an early Paleo- ments (up to 10 km across) of serpentinized harzburgite and lherzolite, zoic granitic pluton, and a subduction–accretion complex containing foliated wehrlite and gabbro (often albitized and amphibolitized), diabase Author's personal copy

68 P. Jian et al. / Earth-Science Reviews 133 (2014) 62–93

Fig. 3. Geological sketch map of the Gobi Altai zone (compiled after GSIGMR, 1998; Hrdličková et al., 2010; Kröner et al., 2010; Lehmann et al., 2010; Štípská et al., 2010). Published zircon data are compiled, and sample localities of this work are indicated.

(dikes) (Zonenshain and Kuzmin, 1978) and minor diorite–tonalite– and 421 ± 3 Ma and 417 ± 2 Ma for two metadacites (Helo et al., trondhjemite (Badarch et al., 2002; Helo et al., 2006), all set in a variably 2006)(Fig. 4). Note that the Ordovician–Silurian volcanic rocks (εNd deformed, greenschist-facies, metavolcanic–sedimentary–volcaniclastic (t) =6–9) and associated siliceous siltstones and volcaniclastic rocks matrix (Lamb and Badarch, 2001; see also Helo et al., 2006; Fig. 4). (6.3–7.4) both have high positive εNd (t) values (Helo et al., 2006), as Lamb and Badarch (2001) suggested that the greenschist facies matrix in the Lake zone (Kovalenko et al., 2004; Kovach et al., 2011). at Zoolen has an approximate Ordovician–Silurian depositional age, that was later corroborated by SHRIMP zircon ages of 481 ± 10 Ma for a 2.1.2.3. South Gobi zone. The South Gobi zone contains a basement meta- meta-tuff (Kröner et al., unpublished; for analytical data see Table S2), morphic complex (Badarch et al., 2002) with orthogneiss (431 ± 2 Ma,

Fig. 4. Geological sketch map of the Gurvan Sayhan and Zoolen areas (compiled after GSIGMR, 1998; Helo et al, 2006). Note that the map is brief: in general, serpentinite is shown, but most of the gabbroic blocks (up to kilometers in size) in the mélanges were not mapped in detail. Published zircon data are compiled, and sample localities of this work are indicated. Author's personal copy

P. Jian et al. / Earth-Science Reviews 133 (2014) 62–93 69

SHRIMP), amphibolite (478 ± 4 Ma, SHRIMP) (Jian et al., 2010b)and Ordovician age (Fig. 3; see also Section 2.1.2.1.3), implying that granodiorite (433 ± 12 Ma; TIMS; Yarmolyuk et al., 2005). The meta- the Gobi Altai ophiolite is not younger than Cambrian–Ordovician. morphic complex is covered by Silurian–Devonian strata (Badarch (3) The matrix of the Gurvan Sayhan–Zoolen mélange has an et al., 2002). Along strike to the east, in the Northern Orogen of Ordovician–Silurian depositional zircon age (Helo et al., 2006; Inner Mongolia of China, there is an early to mid-Paleozoic ridge– Kröner et al., unpublished, for analytical data see Table S2; see subduction complex (Jian et al., 2008). Plutonic and anatectic rocks also Section 2.1.2.3). This implies that the Gurvan Sayhan– of similar early to mid-Paleozoic age also dominate the Chinese Zoolen ophiolite is likely early Paleozoic, like that in the Gobi Altai (Sun et al., 2009)inthefarnorthwest(Fig. 1A), which thus, in our Altai zone. opinion, could be a correlative of the South Gobi zone-Inner Mongolian (4) Along strike in North Xinjiang, northwestern China, the Zhaheba Northern Orogen. ophiolite (SSZ-type) (Fig. 1A) has a well-documented Cambrian– Ordovician age (ca. 503–481 Ma; Jian et al., 2003, 2005; Xiao et al., 2006), which therefore may have formed in the same 2.2. Published zircon data for the ophiolites ocean as the Gurvan Sayhan–Zoolen ophiolite. Plagiogranites in the Dariv and Khantaishir ophiolites have two Taking into account the published geochemical and isotopic signa- conventional, multigrain U–Pb zircon ages. The isotopic composi- tures, the ophiolites in West Mongolia, south of the MML, indicate tions are often discordant, but the weighted mean 207Pb/206Pb or growth of supra-subduction systems in the early Paleozoic. This is con- upper intercept ages are accepted as the time of magma crystalliza- sistent with the fact that many/most ophiolites in the world formed tion: the Dariv ophiolite at 571 ± 4 Ma (Kozakov et al., 2002)and above subduction zones and constitute the basement of oceanic arcs Khantaishir at 568 ± 4 Ma (Gibsher et al., 2001); for analytical data and/or back-arc basins (e.g., Pearce et al., 1984a,b; Shervais, 2001; and interpretations see also Khain et al. (2003).Inspiteoftheir Pearce, 2003; Metcalf and Shervais, 2008; Dilek and Furnes, 2011), discordant isotopic compositions, these zircon ages place a first-order as in the Cenozoic W-Pacific (e.g., Hawkins et al., 1984; Stern and time constraint on the formation of the proto-arc/SSZ-type ophiolites Bloomer, 1992; Bloomer et al., 1995; Pearce et al., 2005) and SW- (e.g., Dijkstra et al., 2006), and hence on subduction initiation in the Pacific(e.g.,Schellart et al., 2006; Whattam et al., 2008; Whattam, central CAOB. 2009; Schellart and Rawlinson, 2010). Considering the fact that it is In South Mongolia ophiolites (Fig. 1A) are widespread in the Gobi widely considered that the crust north of the MML formed in the late Altai zone (Fig. 3) and in the Trans-Altai zone, in the Gurvan Sayhan– Neoproterozoic, it is timely to re-evaluate the significance of this Linea- Zoolen mélanges (Fig. 4). However, the times of formation and em- ment and the consequent tectonic subdivision of West Mongolia, and to placement of these ophiolites are unknown, and there are no reported do that we use the formation/crystallization ages of ophiolites integrat- zircon ages. Thus, the question arises as to when the orogenic activity ed with other published data. began in South Mongolia, and this reveals a fundamental problem with the regional tectonic subdivision, as discussed below. 3. Approach to constrain the formation of ophiolites 2.3. A fundamental problem in tectonic subdivision The dating of ophiolites is generally based on the isotopic systems of West Mongolia has long been tectonically divided into two do- magmatic crustal rocks. Mantle peridotites bear little or no geochrono- mains by the MML (Fig. 1A), namely early Paleozoic (formerly called logical significance, because they mostly are not magmatic products but Caledonian) in the north (i.e. Northwest Mongolia of this paper) and late melt-residuals after basaltic magma extraction (e.g., Rampone et al., Paleozoic (formerly called Hercynian) in the south (i.e. South Mongolia) 1996). The magmatic crustal rocks, on the other hand, are relatively uni- (e.g., Amantov et al., 1970; for overviews see Badarch et al., 2002; form in composition, are usually altered to greenschist-facies, and con- Kröner et al., 2010; Lehmann et al., 2010). From recently published geo- tain few or no zircons, and are often unsuitable for Sm–Nd, Rb–Sr, and logical and geochronological data, we note a fundamental problem in K–Ar, Ar–Ar or conventional multigrain zircon dating. However, zircon the previous subdivision, in particular with the late Paleozoic belts of dating using a high-resolution ion microprobe or high-precision TIMS South Mongolia. (Thermal Ionization Mass Spectrometer) needs only a few zircons that The late Paleozoic belts were defined geochemically by Permo- can often be routinely obtained from ophiolitic gabbros, plagiogranites Carboniferous volcanic rocks and granitoids (e.g., Lamb and Badarch, and intermediate to felsic rocks (e.g., Schwartz et al., 2005; Baines 2001; Yarmolyuk et al., 2008b) and in terms of ubiquitous negative et al., 2009; Grimes et al., 2009; see also Jian et al., 2010a,b). Nb–Ta anomalies. Unfortunately, the volcanic rocks (e.g. trachyandesite, trachyrhyolite) and granitoids (e.g., monzonite, syenite) are mostly 3.1. Appropriateness of gabbroic and leucocratic rocks to zircon chronology alkaline and are therefore, in a rather general sense, petrologically inconsistent with volcanic (tholeiitic and calc-alkaline; basalt–andesite– 3.1.1. Gabbroic rocks dacite–rhyolite; e.g., Bailey, 1981; Macdonald et al., 2000)andplutonic In the classic ophiolite stratigraphy (Anon, 1972), there are two (calc-alkaline and sodic; diorite–quartzdiorite–tonalite–trondhjemite; distinct structural types of gabbro: a layered type representing the e.g., Barton et al., 1988; Martin et al., 2005) arcs. We thus question this lower oceanic crust, and an isotropic gabbro that formed in the mid- interpretation. Instead, currently available evidence strongly argues upper oceanic crust (Nicolas, 1989; Boudier et al., 1996; Kelemen and for early Paleozoic orogenic terranes in South Mongolia, as outlined in Aharonov, 1998). The formation of a layered gabbro in an ophiolite the following: has been ascribed to crystal settling in a magma chamber beneath a (1) An angular unconformity in the Gobi–Altai zone separates a spreading ridge (Browning, 1984; Kelemen et al., 1997; for overview fossil-dated, early Ordovician conglomerate from underlying see also Coogan, 2007). Anorthosite, a gabbroic variant that consists of schists (Kröner et al., 2010). The unconformity postdates ca. more than 90% plagioclase, often occurs as discontinuous, thin (up 518–498 Ma anatectic granitic rocks (Hrdličková et al., 2008; to several tens of centimeters) layers and lenses in layered gabbros Demoux et al., 2009b; Hrdličková et al., 2010; Kröner et al., (e.g., Hoeck et al., 2002). The gabbroic rocks, including the layered and 2010) (see also Section 2.1.2.1.1), which implies a period of isotropic types and anorthosite all contain zircons, largely depending Cambrian orogenesis. on the modal content of plagioclase, and are therefore favorable for (2) The metavolcanic–sedimentary matrix of a mélange that con- SHRIMP zircon dating. Most zircons in modern oceanic (e.g., Schwartz tains numerous ophiolitic fragments has a possible Cambrian– et al., 2005; Baines et al., 2009; Grimes et al., 2009) and ancient Author's personal copy

70 P. Jian et al. / Earth-Science Reviews 133 (2014) 62–93 ophiolitic (e.g., Kaczmarek et al., 2008) gabbros formed by magmatic 4. Field investigation and samples crystallization. 4.1. Dariv and Khantaishir (Northwest Mongolia) 3.1.2. Leucocratic rocks (oceanic plagiogranites) 4.1.1. Dariv Plagiogranites typically occur as dikes and pods in the upper parts The ophiolite in the northwestern Dariv Range has a near-complete of ophiolites (Coleman and Peterman, 1975), and lithologically include magmatic stratigraphy (Fig. 2A), which is ideal for SHRIMP zircon dating albite granite, quartz diorite, trondhjemite, tonalite, keratophyre and and relevant geochemical studies. We collected three samples for dating albitite. These rocks that are chiefly composed of quartz and plagioclase from a sheeted dike complex (Fig. 2A): a coarse-grained plagiogranite with only minor ferromagnesian minerals and share several com- (M99D11), a fine-grained plagiogranite (MDRV09-3; Fig. 5A) and a positional features: high silica (Si), moderate alumina (Al), low microgabbro (MDRV09-1; Fig. 5B). The sheeted dike complex comprises iron-magnesium (Fe–Mg) and extremely low potassium (K). In some numerous meter-scale mutually parallel dikes of predominantly basal- classic ophiolites, like the Semail in Oman (e.g., Nicolas, 1989) and the tic composition (i.e. microgabbros and diabases). Note that at the out- Troodos in Cyprus (e.g., Robinson and Malpas, 1990), the plagiogranites crop of sample MDRV09-3, a diorite dike (e.g., MDRV09-2) transects may form intrusive plugs and parallel dikes within diabase dike the microgabbro at a high angle (e.g., Fig. 5B). One plagiogranite is a swarms. Oceanic plagiogranites are most commonly explained by pro- coarse-grained (M99D11), narrow dike (ca. 5 cm wide) that cuts gressive differentiation of a tholeiitic melt (e.g., Coleman and microgabbro, the other forms a fine-grained (MDRV09-3) small patch Peterman, 1975; Pearce et al., 1984b; Toplis and Carroll, 1995) though (b0.25 m) in microgabbro (Fig. 5A). The plagiogranites are composed other petrogenetic processes are possible (e.g., Koepke et al., 2004), of predominant plagioclase (N60 vol.%), subordinate quartz (bca. and their crystallization ages reflect ophiolite formation through 30%), and minor, secondary chlorite, epidote, calcite and muscovite, sea-floor spreading (e.g., Robinson et al., 2008). and accessory minerals. Microgabbro sample MDRV09-1 is massive, fine-grained and has a poikilitic texture. It has nearly equal contents of 3.2. Methodology clinopyroxene and plagioclase, along with secondary chlorite and epidote; also notably some fine-grained felsic veins. The age of an ophiolite can thus be constrained via single zircon Three additional samples were collected from the adjacent dating of gabbroic and leucocratic rocks, as stated above. However, in gabbro–pyroxenite complex in the south (Fig. 2A). Samples MDRV07-1 some ophiolite mélanges, there may be overprinted, post-ophiolite, and MDRV07-2 are coarse-grained, isotropic olivine gabbros that magmatic products of similar gabbroic (e.g., Jian et al., 2010a)and contain ca. 3–5% olivine, ca. 35–40% clinopyroxene and more than leucocratic (e.g., Shervais, 2008) lithologies. In order to overcome this 50% plagioclase, with secondary chlorite and epidote, but unfortu- potential confusion, we identify true ophiolitic gabbroic and leucocratic nately, without significant zircons. Sample M99D19 is from an unde- rocks by their geochemical signatures, supplementary to the field inves- formed hornblende granite dike (ca. 0.2 m wide) that cuts an isotropic tigation and zircon chronology. gabbro. In general, gabbroic rocks contain less zircon than leucocratic rocks. In order to extract representative zircon populations in the field we col- 4.1.2. Khantaishir lected gabbroic samples, each about 20–50 kg, and leucocratic samples, The Khantaishir ophiolite forms tectonic slices of harzburgite– each ca. 5–10 kg. U–Th–Pb zircon analyses were performed by SHRIMP dunite, gabbro–pyroxenite, sheeted dikes and lavas, and widespread II at the Institute of Geology, Chinese Academy of Geological Sciences, serpentinite mélanges (Fig. 2B). We collected two plagiogranite Beijing; the analytical procedures are summarized elsewhere (e.g., Jian samples to date the formation of the ophiolite. Sample M98K14 is a et al., 2010a,b). Errors of individual analyses, given at the 1-sigma loose block (ca. 1 m long) in a serpentinite mélange (Fig. 2B). Sample level in Supplementary Table S1 (this work) and Table S2 (Kröner M98K12 is a thin dike (ca. 5 cm wide; Fig. 5C) in gabbro from a gabbro– et al., unpublished), were determined from counting statistics and pyroxenite slice (for location see Fig. 2B). These two plagiogranites are time-dependent correction of the standard (i.e. TEM; Black et al., fresh, coarse-grained and contain predominant plagioclase (ca. 65–70%), 2003) calibration constant (Compston, 1999). This 1-sigma error was subordinate quartz (25–30%), and minor amounts of secondary chlorite used to construct error boxes for individual data points on all concordia and epidote. plots. Unless otherwise stated, all magmatic/metamorphic ages report- As shown in Fig. 2B, ca. 10 km west of Altai City, a serpentinite mé- ed in the text and summarized in Table 1 are weighted mean 206Pb/238U lange is thrust as a tectonic slice onto the Dzabkhan microcontinent. ages, and uncertainties in the figures and in the text are cited at 95% In the mélange (i.e. an ophiolite slice) there are numerous tectonic blocks, confidence limit. up to several hundred meters across, of pillow lava (e.g., Fig. 5D) of Major and trace elements of most samples were analyzed in the basaltic (e.g., sample MHTS01) to andesitic (e.g., samples MTH06 and Laboratory of Continental Dynamics, Northwest University, Xian, MTHS04) composition as well as gabbros (e.g., sample MHTS02-2) and China, using standard X-ray fluorescence (XRF) techniques and an pyroxenites, all set in a dominant, strongly sheared serpentinite matrix. ICP-MS with Elan 6100DRC plasma. Other analytical work was under- On the boundary between the serpentinite mélange (ophiolite) and the taken at the National Research Center of Geoanalysis, China, using an microcontinent, we collected a strongly foliated amphibolite sample XRF spectrometer for major element analysis, and an ICPMS (X-series) (MHTS07-1; Fig. 2B), which consists of alternating amphibole-rich and for trace elements. Analytical uncertainties were ~1.5% for major ele- plagioclase-rich layers (Fig. 5E). Nearby, there is garnet gneiss in the ments, and ~5–15% for trace elements, depending on concentration microcontinent. The amphibolite most likely records the time of meta- level. morphism that accompanied an important tectonic movement, and the Nd and Sr isotopes on selected whole-rock samples were deter- igneous rocks from the mélange, in particular the lavas, are the best mined in the Institute of Geology, Chinese Academy of Geological geochemical indicators of the original tectonic setting of the ophiolite, Sciences, Beijing, using conventional column chemistry and a thermal and so we investigate them as part of this study. ionization mass-spectrometer (MAT 261). 143Nd/144Nd ratios were corrected for mass fractionation relative to 146Nd/144Nd = 0.7219 and 4.2. Gobi Altai (South Mongolia) are reported with reference to the La Jolla Nd standard = 0.511860. Mass fractionation correction for 87Sr/86Sr is relative to 86Sr/88Sr = The Gobi Altai ophiolite is tectonically dismembered, and is partially 87 86 0.1194. Initial εNd values and Sr/ Sr ratios were calculated for mag- thrust on a Precambrian microcontinent and partially associated with a matic ages as determined by zircon analyses. volcano-sedimentary mélange of possible Cambrian–Ordovician age Author's personal copy

Table 1 Summary of geological, geochemical and SHRIMP zircon ages.

87 86 Locality Sample no. Lithology Field occurrence Location Age εNd(t) Initial Sr/ Sr (Ma)

Lat. (N) Long. (E) Magmatic Metamorphic Xenocrysts (inherited zircons) .Ja ta./ErhSineRves13(04 62 (2014) 133 Reviews Earth-Science / al. et Jian P. Dariv M99D11 Plagiogranite (coarse-grained) Dike within sheeted dike 46°48.069′ 94°02.161′ 560 ± 8 Ma; n = 15; χ2 =0.49 complex MDRV09-3 Plagiogranite (ﬁne-grained) Small patch in microgabbro ca. 500 m north of MDRV09-1 567 ± 4 Ma; n = 16; χ2 = 1.02 5.3 0.7044 within sheeted dike complex MDRV09-1 Microgabbro Dike within sheeted dyke 46°42′ 33.1″ 94°08′ 13.6″ 566 ± 5 Ma; n = 12; χ2 =0.61 complex M99D19 Leucogranite Dike cutting isotropic 46°37.760′ 94°07.500′ 487 ± 7 Ma; n = 13; χ2 = 0.51 ca. 745–764 Ma −11.7 0.7086 gabbro Khantaishir M98K12 Plagiogranite Dike in isotropic gabbro 46°18.096′ 96°10.379′ 565 ± 7 Ma; n = 18; χ2 = 1.01 6.6 0.7037 M98K14 Small block in serpentinite ca. 200 m south of M98K12 573 ± 8 Ma; n = 12; χ2 =0.51 MHTS07-1 Amphibolite A shear zone on the 46°25′ 43.8″ 96°06′ 30.7″ 514 ± 8 Ma; n = 12; ca. 608–2433 Ma ophiolite/microcontinent χ2 = 0.29 boundary Gobi Altai MALT18-1 Gabbro (layered) Tectonic block in serpentine 45°28′ 22.4″ 98°02′ 42.4″ 523 ± 5 Ma; n = 10; χ2 = 2.12 7.4 0.7028 ′ ″ ′ ″ χ2

MALT14 Leucogabbro (isotropic) Tectonic block in sediments 45°5 21.3 99°29 3.3 518 ± 6 Ma; n = 14; = 0.16 4.3 0.7048 – Gurvan Sayhan ZL03-1 Hornblendite Tectonic block in sediments 43° 50′ 01.5″ 103° 02′ 52.7″ ≥512 ± 4 Ma; n = 7; χ2 = 1.85 7.3 0.7038 93 ZL03-2 Anorhosite Small patch (0.2 × 0.5 m) in 519 ± 4 Ma; n = 9; χ2 = 0.61 7.9 0.7044 hornblendite Zoolen ZL09-1 Diorite Tectonic block in sediments 43° 32′58″ 102° 42′8.1″ 511 ± 5 Ma; n = 13; χ2 = 0.39 7.3 0.7039 ZL10-1 Diorite Tectonic block in sediments 43°37′19.2″ 102° 32′ 21.1″ 520 ± 5 Ma; n = 15; χ2 =1.24 ZL10-2 Tonalite Dike cutting diorite 449 ± 6 Ma; n = 8; χ2 =1.46 ZL12-1 Tonalite Dike cutting a serpentinite 43°35.985′ 102° 37.660′ 494 ± 6 Ma; n = 8; χ2 = 0.41 7.7 0.7045 melange 71 Author's personal copy

72 P. Jian et al. / Earth-Science Reviews 133 (2014) 62–93

A B

C D

E F

G H

Fig. 5. Photographs showing ﬁeld relationships of important samples. (A) Plagiogranite (MDRV09-3) in microgabbro, Dariv; (B) Microgabbro (MDRV09-1) and diorite (MDRV09-2) dikes within a sheeted dike complex, Dariv; (C) Plagiogranite (white dike) in gabbro, Khantaishir; (D) Pillow lava (MHTS06; andesite) in a serpentinite mélange, Khantaishir; (E) Strongly foliated amphibolite (MHTS07-1) in a thrust/shear zone between the Khantaishir ophiolite and the Dzabkhan–Baydrag microcontinent; (F) Leucogabbro (MALT14), Gobi Altai; (G) Gabbro blocks in serpentinite, Gurvan Sayhan; (H) Anorthosite (ZL03-2) occurring as a patch in foliated hornblendite (ZL03-1), Gurvan Sayhan.

(Fig. 3). We investigated a large ophiolite slice (ca. 15 km long, 3 km altered), with minor Fe–Ti oxides and secondary epidote, chlotite, wide) on the microcontinent north of Chandman (Fig. 3). This slice and calcite. consists predominantly of serpentinite (strongly sheared), in which In order to establish the age relationships between the ophiolite dunites, pyroxenites (usually b10 m in size) and gabbros (up to several fragments in slices thrust onto the microcontinent and those in the hundred meters across) occur as tectonic blocks. Zircon-dated gabbro volcano-sedimentary mélanges, we collected a leucogabbro sample (sample MALT18-1; Fig. 3) was from a relatively large, ca. 300 m (MALT14; Fig. 3) from the latter. The leucogabbro forms a small wide block that is cut sharply by several fresh diabase (MALT18-2) (ca. 10 m long), isolated fragment within a volcano-sedimentary matrix, and felsic dikes up to 2 m wide. Sample MALT18-1 contains plagioclase- and has no direct contact with serpentinite. However, nearby, there is a rich and pyroxene-rich magmatic layers, and almost equal amounts ca. 300 m long serpentinite block. As the layered gabbro (MALT18-1), of clinopyroxene (essentially primary) and plagioclase (partially leucogabbro sample MALT14 consists of clinopyroxene (essentially Author's personal copy

P. Jian et al. / Earth-Science Reviews 133 (2014) 62–93 73 primary) and plagioclase (partially altered), though it is strongly foliat- Fine-grained plagiogranite sample MDRV09-3 contains subhedral ed as shown in Fig. 5F. zircon grains, which, as in the coarse-grained plagiogranite, have concordant isotopic compositions that again form a single cluster (χ2 = 1.02; 4.3. Gurvan Sayhan and Zoolen (Trans-Altai zone, South Mongolia) Fig. 7B). Note that, as indicated by the relatively weak CL-emissions (Fig. 6-2), their U concentrations (125–1111 ppm; mostly N 400 ppm; Although the Gurvan Sayhan and Zoolen ophiolites are poorly known, Table S1) are distinctively high when compared with those of the they are represented by serpentinite mélanges (e.g., Zonenshain, 1973) coarse-grained plagiogranite (ca. 77–264 ppm). This results in a higher and gabbroic–dioritic blocks (e.g., Badarch et al., 2002; Helo et al., concentration level of radiogenic lead of zircons (Table S1) and thus an 2006). The presently available map (Fig. 4) shows serpentinites but not overall better analytical precision (i.e. smaller error boxes) than that of the gabbroic–dioritic blocks, which are the subject of our current detailed the coarse-grained plagiogranite. The weighted mean age is 567 ± study. 4 Ma (n = 16; Fig. 7B) that we interpret as the formation age of this plagiogranite. 4.3.1. Gurvan Sayhan The Gurvan Sayhan serpentinite (samples ZL01-2 and ZL02-2) locally 5.1.1.2. Microgabbro of the sheeted dike complex. Most zircon grains of contains strongly sheared gabbroic blocks (e.g., Fig. 5G), which unfortu- microgabbro sample MDRV09-1 are small (ca. 50 μm), anhedral and nately do not contain sufficient and representative zircon grains to strikingly weak in CL-emissions (e.g., Fig. 6-3, left). Their analyses measure by SHRIMP. Instead we collected from the same outcrop an yielded a concordant cluster with a weighted mean age of 568 ± 5 Ma anorthosite (ZL03-2) and a hornblendite (ZL03-1) from a tectonic block (n = 12, χ2 = 0.61; Fig. 7C). This age is identical, within error, to the (ca. 1 km2 in areal content) of hornblendite in a volcaniclastic matrix above reported 567 ± 4 Ma age for the associated, fine-grained (tuff + chert) (Fig. 4). In the outcrop, anorthosite occurs as a small plagiogranite (sample MDRV09-3), and is thus interpreted as the time patch (ca. 20 cm) in a strongly foliated hornblendite (Fig. 5H). Anorthosite of microgabbro crystallization. sample ZL03-2 is coarse-grained and contains predominant plagioclase The remaining five zircon grains of the same sample yielded ages of (ca. 85%), subordinate amphibole (ca. 15%), and minor secondary epidote ca. 271–459 Ma (Table S1; Fig. 7C), much younger than the above for- (ca. 5%). Hornblendite sample ZL03-1 consists of amphibole (ca. 60%), mation age of the microgabbro. These zircons contain fine oscillatory plagioclase (ca. 25%), and minor epidote and chlorite (each ca. 5%). and concentric growth zones (e.g., Fig. 6-3, right), indicating crystallization from a felsic magma (e.g., Hanchar and Miller, 1993). We propose 4.3.2. Zoolen that such “magma” is represented by the thermally-generated fine- We sampled two diorites (samples ZL10-1 and ZL09-1; Fig. 4) grained felsic veins, which are pervasive in the dated microgabbro sam- from Zoolen, which form two elliptical tectonic blocks, each of ca. ple (see Section 4.1.1). 100 m2, one in metavolcanic rocks (ZL10-1) the other (ZL09-1) in metavolcaniclastic rocks. They are weakly foliated and are cut by un- 5.1.1.3. Leucogranite dike cutting isotropic gabbro. Most zircons separated deformed dikes (b10 cm wide) of tonalite (sample ZL10-2) and dia- from leucogranite (dike) sample M99D19 are large (N100 μm) (origi- base (sample ZL09-2). Diorite sample ZL10-1 is medium-grained and nally) euhedral crystals with oscillatory zoning (e.g., Fig. 6-4, bottom comprises amphibole (ca. 50%), plagioclase (ca. 45%), Fe–Ti oxides two). Their concordant ages are grouped at 487 ± 7 Ma (n = 13, (b1%) and secondary epidote, chlorite and calcite (b4%). Under the χ2 =0.51;Fig. 7E), which we interpret as reflecting the emplace- microscope, microveins (epidote + feldspar) are visible. Diorite sample ment time of the leucogranite dike. In addition, we measured two ZL09-2 has a similar mineral composition, but is coarse-grained and xenocrystic zircons, which have older ages of ca. 745 and ca. 764 Ma contains megacrysts of feldspar and amphibole. (207Pb/206Pb ages; Table S1, Fig. 7E), broadly similar to the ages of the Finally, a serpentinite (sample ZL12-2) from a mélange contains mid-Neoproterozoic diorites–granites in the Dzabkhan–Baydrag block variably-sized tectonic blocks of gabbro, basalt and sandstone, and is (ca. 755–856 Ma; Zhao et al., 2006; Yarmolyuk et al., 2008a; Levashova intruded by a ca. 2 m wide, relatively large tonalite dike. The tonalite et al., 2010; Kozakov et al., 2012). Interestingly, these two xenocrysts (sample ZL12-1) is medium-grained, weakly foliated and contains are essentially euhedral crystals, either long- (Fig. 6-4, right middle) plagioclase (ca. 60%), quartz (ca. 20%), Fe–Ti oxides (ca. 1.5%), with sec- or short-prismatic (Fig. 6-4, right top), without morphological indica- ondary epidote and chlorite (ca. 18% in total). tion for long-distance transport.

5. Zircon ages 5.1.2. Khantaishir

The internal textures of the zircons were studied by cathodo- 5.1.2.1. Plagiogranites of the ophiolite. Zircons of the two plagiogranite luminescence (CL) imaging, and representative images are presented samples are from a serpentinite mélange (M98K14), and from a in Fig. 6. In general, weak CL-emissions (dark areas in the images) cor- gabbro–pyroxenite complex (M99K12), both are prismatic, either oscil- respond to high U, Th and other trace elements, and strong CL emissions latory or patchily zoned, original magmatic crystals (Figs. 6-5, 6-6). (light areas) reflect the opposite (Hanchar and Miller, 1993). Their crystallization ages, as determined by 12 and 18 analyses respectively, are 573 ± 8 Ma (n = 12, χ2 =0.51;Fig. 7E) and 566 ± 7 Ma 5.1. Dariv and Khantaishir (Northwest Mongolia) (n = 18, χ2 = 1.01; Fig. 7F), identical within analytical error. These ages are interpreted to represent the time of plagiogranite formation. 5.1.1. Dariv 5.1.2.2. Amphibolite on the ophiolite/microcontinent boundary. Most 5.1.1.1. Plagiogranites in a sheeted dike complex. Zircons from coarse- zircons in amphibolite sample MHTS07-1 have a core-rim structure grained plagiogranite (sample M99D11) are uniformly prismatic crys- (Fig. 6-7), common in high-grade metamorphic rocks (e.g., tals (ca. 50–100 μm in size). Their CL images (Fig. 6-1) show patchy Williams et al., 1996). The cores yielded normal discordant isotopic and oscillatory growth zones, characteristic of magmatic crystallization compositions and a wide range of 207Pb/206Pb ages varying between in ancient and modern oceanic plagiogranites (Schwartz et al., 2005; ca. 608 Ma and 2433 Ma (Table S1; Fig. 7G). Most likely, the amphib- Baines et al., 2009; Grimes et al., 2009). The isotopic compositions plot olite has a sedimentary protolith, and the zircon cores reflect a complex as a single concordant group with a weighted mean 206Pb/238U age of detrital source. 560 ± 8 Ma (n = 15, χ2 = 0.49) (Fig. 7A), which we interpret to reflect The rims are mostly thin and discontinuous (e.g., Fig. 6-7, right two) the time of (zircon and) plagiogranite crystallization. and only locally wide enough (≥25–30 μm; e.g., Fig. 6-7, left three) Author's personal copy

74 P. Jian et al. / Earth-Science Reviews 133 (2014) 62–93

Fig. 6. Representative CL images of zircons. 1—M99D11 (plagiogranite, Dariv); 2—MDRV09-3 (plagiogranite, Dariv); 3—MDRV09-1 (microgabbro, Dariv); 4—M99D19 (hornblende granite, Dariv); 5—M98K14 (plagiogranite, Khantaishir); 6—M98K12 (plagiogranite, Khantaishir); 7—MHTS07-1 (amphibolite, Khantaishir); 8—ZL03-2 (anorthosite, Gurvan Sayhan); 9—ZL03-1 (hornblendite, Gurvan Sayhan); 10—ZL10-1 (diorite, Zoolen); 11—ZL09-1 (diorite, Zoolen); 12—ZL12-1 (tonalite, Zoolen); 13—ZL10-2 (tonalite, Zoolen). Author's personal copy

P. Jian et al. / Earth-Science Reviews 133 (2014) 62–93 75

Fig. 7. Concordia diagrams for zircons of samples from the Dariv (A–D)–Khantaishir (E–H) ophiolites and related lithological units. A—Plagiogranite (M99D11); B—Plagiogranite (MDRV09-3); C—Microgabbro (MDRV09-1); D—Hornblende granite (M99D19); E—Plagiogranite (M98K14); F—Plagiogranite (M98K12); G—Amphibolite (MHTS07-1; all data); H—Amphibolite (MHTS07-1; rims). Data boxes are deﬁned by standard errors (1 sigma) in 206Pb/238U, 207Pb/235Uand207Pb/206Pb. for SHRIMP measurement. The measured rims have low U contents consistent with a metamorphic origin (e.g., Williams et al., 1996). The (4–48 ppm; Table S1) and extremely low Th contents (b1–8 ppm, metamorphic age of this rock is thus determined at 514 ± 8 Ma (χ2 = Table S1) that result in low Th/U ratios (mostly, b0.1), which are 0.29; Fig. 7G and H) by 12 concordant analyses. Author's personal copy

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Fig. 8. Concordia diagrams (left) and CL images (right) of zircons from samples of the Gobi–Altai ophiolite belt. A and B—MALT18-1 (gabbro); C and D—MALT14 (leucogabbro). Data boxes as in Fig. 7.

5.1.3. Summary (Khantaishir) is dated at 566 ± 7 Ma, indistinguishable from the forma- The above results (Table 1)conﬁrm that the Dariv–Khantaishir tion age of a sheeted dike complex (Dariv)—in the upper part of the ophiolite formed at ca. 570 Ma (Gibsher et al., 2001; Kozakov et al., ophiolite; (3) the formation age of the sheeted dike complex is deter- 2002; Khain et al., 2003). Looking into the SHRIMP zircon data mined by the identical ages of a microgabbro dike (566 ± 5 Ma) and a (Table 1) we address the following points: (1) the oldest-known ﬁne-grained plagiogranite patch (567 ± 4 Ma); and (4) the youngest, plagiogranite (tectonic block; 573 ± 8 Ma; Table 1) is associated with coarse-grained plagiogranite (560 ± 8 Ma) occurs a small cross-cutting a serpentinite mélange (Khantaishir)—at the tectonically lowest dike in the sheeted dike complex. These possibly reveal the relative level; (2) a plagiogranite dike from a gabbro–proxenite complex timing as well as the subtle age differences between the ophiolite units. Author's personal copy

P. Jian et al. / Earth-Science Reviews 133 (2014) 62–93 77

C D

E F

Fig. 9. Concordia diagrams for zircons from samples of the Gurvan Sayhan (A and B) and Zoolen (C–F) sedimentary and ophiolitic mélanges. A—Anorthosite (ZL03-2); B—Hornblendite (ZL03-1); C—Diorite (ZL10-1); D—Tonalite (ZL10-2); E—Diorite (ZL09-1); F—Tonalite (ZL12-1). Data boxes as in Fig. 7.

The relative timing is consistent with the idealized ophiolite sequence relatively large grains (ca. 40–100 μm; Fig. 8C). These grains mostly dis- (e.g., Nicolas, 1989; Boudier et al., 1996; Kelemen and Aharonov, 1998). play patchy zoning, like that of magmatic zircons of ancient ophiolitic The age difference between the oldest (at a tectonically low level) and (e.g., Kaczmarek et al., 2008) and modern oceanic (Schwartz et al., the youngest (at a tectonically high level) plagiogranites is about 2005; Baines et al., 2009; Grimes et al., 2009) gabbros. The weighted 13 Ma, which we interpret to approximate the duration of the ophiolitic mean age of all analyses is 523 ± 5 Ma (n = 10; Fig. 8A), which we crust formation that is as short as normal (for overview see Hopson et al., suggest approximates the formation time of this layered gabbro; 2008). The formation time of the Dariv–Khantaishir ophiolite can thus be although a relatively large χ2 of 2.12 (N1) indicates some excess scatter constrained to the period ca. 573–560 Ma. Additional age data for the of the age population (c.f. Black and Jagodzinski, 2003). We prefer an amphibolite (ca. 514 Ma metamorphism; Khantaishir; Table 1)on arbitrary weighted mean calculation (without rejects) simply because the ophiolite/microcontinent boundary and the leucogranite dike of a limited dataset. (ca. 487 Ma granitoid magmatism; Dariv) cutting isotropic gabbro place time constraints on the post-ophiolite evolution, and their signiﬁcances will be explored in some detail in the Discussion section 5.2.2. Isotropic leucogabbro from a southern volcano-sedimentary mélange (Section 7). We extracted 12 zircon grains (Fig. 8D) from leucogabbro sample MALT14 and measured fourteen spots on these. As in the layered 5.2. Gobi Altai (South Mongolia) gabbro, the zircons of this sample are all gabbroic magmatic crystals judging from their CL patterns (Fig. 8D). Their isotopic compositions 5.2.1. Layered gabbro from a northern thrust slice plot as a concordant cluster with a weighted mean age of 518 ± 6 Ma Layered gabbro sample MALT18-1 contained scarce zircons, some of (n = 14; χ2 = 0.16; Fig. 8B), which best dates the time of formation which were too small (b30 μm in width) to measure, but we dated 10 of the leucogabbro. Author's personal copy

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Table 2 Major (wt.%), REE and trace element (ppm) compositions of samples.

Locality Dariv

Unit Sheeted dike complex Gabbro–pyroxenite

Sample no. MDRV09-1 MDRV09-2 MDRV09-3 M99D19a MDRV07-1 MDRV07-2

Lithology Microgabbro Diorite (dike) Plagiogranite Leucogranite Olivine gabbro (dike; 568 ± 5 Ma) (567 ± 4 Ma) (dike; 487 ± 7 Ma)

Oxides (%)

SiO2 47.04 54.19 71.24 72.67 47.03 47.19

TiO2 0.07 0.23 0.12 0.51 0.11 0.1

Al2O3 18.65 12.87 12.59 13.24 12.68 10.7 T Fe2O3 6.34 7.92 2.4 6.69 8.12

Fe2O3 0.68 FeO 2.23 MnO 0.12 0.1 0.05 0.1 0.13 0.16 MgO 8.82 6.61 1.06 1.12 14.77 16.19 CaO 7.02 7.9 5.84 1.46 12.25 12.09

Na2O 2.87 2.55 4.4 3.58 0.89 1.06

K2O 1.54 0.21 0.17 3.07 1.28 0.63

P2O5 0.01 0.05 0.02 0.11 0.01 0.01

H2O 1.01

CO2 0.52 LOI 7.69 6.91 2.41 1.05 3.68 3.39 Total 100.17 99.54 100.3 101.35 99.52 99.64

REE and trace element (ppm) La 0.88 5.25 3.67 39.8 0.39 0.73 Ce 1.94 10.9 6.97 82.5 0.88 1.27 Pr 0.24 1.29 0.75 9.19 0.13 0.15 Nd 1.09 5.91 3.27 35.4 0.73 0.73 Sm 0.33 1.55 0.84 7.13 0.29 0.26 Eu 0.14 0.81 0.37 1.52 0.14 0.12 Gd 0.4 1.72 1.03 6.98 0.43 0.37 Tb 0.072 0.28 0.18 1.01 0.081 0.071 Dy 0.49 1.8 1.27 6.23 0.56 0.5 Ho 0.12 0.4 0.32 1.39 0.13 0.12 Er 0.35 1.16 1.02 4.13 0.38 0.34 Tm 0.055 0.18 0.18 0.62 0.06 0.053 Yb 0.38 1.22 1.3 4.18 0.4 0.35 Lu 0.057 0.19 0.22 0.65 0.06 0.054 Y 2.73 10.5 9.14 33.7 2.99 2.87 Sc 28.2 27.8 5.67 10.2 33.3 91.5 V 122 233 57.5 37.5 164 167 Cr 164 391 28.8 17.1 793 791 Co 33.5 36.3 6.03 5.4 45.5 60.4 Ni 73.3 70 14.9 8.52 163 262 Cu 15.2 12.9 11.8 28.8 24.7 72.2 Zn 29.1 26.2 6.84 53.8 35.6 38.5 Ga 12.7 15.8 9.65 13.9 7.73 7.34 Rb 39.5 3.22 2.78 52 40.1 18.2 Sr 331 275 185 364 271 177 Zr 6.23 34.9 83.1 255 1.69 1.22 Nb 0.24 0.76 0.76 8.44 0.082 0.055 Ta 0.022 0.061 0.076 0.58 0.011 0.01 Cs 3.95 0.49 0.11 1.26 3.07 0.56 Ba 142 125 45.2 1081 136 56.3 Hf 0.21 1.08 2.75 7.23 0.092 0.073 Pb 1.56 2.16 1.92 53.3 2.33 2.13 Th 0.24 1.03 2.6 10.6 0.029 0.078 U 0.093 0.56 0.46 1.7 0.01 0.019

a Analyzed at the National Research Center of Geoanalysis, China.

5.2.3. Summary 5.3. Gurvan Sayhan and Zoolen (Trans-Altai, South Mongolia) The results demonstrate that the two gabbros, i.e. the layered gabbro (523 ± 5 Ma; Table 1) from the northern thrust slice and the isotropic 5.3.1. Anorthosite (patch) and hornblendite (tectonic block) in a volcano- leucogabbro (518 ± 6 Ma) from the southern volcano-sedimentary mé- sedimentary mélange (Gurvan Sayhan) lange, have similar formation ages. The minor age difference (ca. 5 Ma) The anorthosite (sample ZL03-2) contains prismatic (either long or between the two structural types, i.e., layered versus isotropic, is readily short), euhedral zircon crystals that are oscillatory and patchily zoned explained by the plutonic sequence of a classic ophiolite stratigraphy (e.g., Fig. 6-8). Nine analyses (Table S1) are concordantly clustered at (Nicolas, 1989; Boudier et al., 1996; Kelemen and Aharonov, 1998). 519 ± 4 Ma (χ2 = 0.61) (Fig. 9A), which is interpreted as the crystalli- This helps to reconstruct the ophiolite stratigraphy, and we accept the zation age of the anorthosite. ca. 523–518 Ma age range as reﬂecting the formation time of the Gobi More than half of the zircons from hornblendite sample ZL03-1 are Altai ophiolite. relatively euhedral, either long- or short-prismatic crystals that are Author's personal copy

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Table 2 Major (wt.%), REE and trace element (ppm) compositions of samples.

Khantaishir Gobi Altai Gurvan Sayhan

Lava Gabbro–pyroxenite A shear zone Gabbro massif Gabbro layer above serpentinite Serpentinite mélange

MHTS01 MHTS06 MHTS04 M98K12a MHTS02-2 MHTS07-1 MALT18-1 MALT18-2 MALT14 ZL01-2 ZL02-1

Andesitic Andesite–dacite Plagiogranite Olivine gabbro Amphibolite Gabbro Diabase (dike) Gabbro Serpentinite basalt (566 ± 7 Ma) (514 ± 8 Ma) (523 ± 5 Ma) (518 ± 6 Ma) (metaharzburgite–dunite)

Oxides (%) 54.24 57.55 63.08 77.47 48.29 49.81 44.37 47.7 51.13 35.72 39.71 0.37 0.59 0.38 0.33 0.1 0.76 0.16 1.11 0.17 b0.01 b0.01 14.48 15.47 13.71 11.76 18.09 14.11 23.61 18.39 16.78 0.89 0.67 9.34 8.22 8 8.49 11.63 4.79 9.01 4.98 9.45 6.67 0.042 0.84 0.13 0.14 0.1 0.018 0.16 0.17 0.09 0.14 0.1 0.13 0.11 6.46 5.66 5.53 1.07 8.08 7.63 7.44 6.97 7.59 35.01 39.54 4.17 3.7 1.77 2.52 12.48 11.19 12.27 9.6 12.32 4.06 0.32 3.73 4.22 4.71 4.55 0.53 2.87 2.58 3.86 3.59 b0.01 b0.01 0.77 1.37 0.1 0.23 0.05 0.35 0.18 0.43 0.1 0.02 0.01 0.05 0.07 0.04 0.045 0.01 0.06 0.01 0.29 0.02 0.01 0.01 0.74 0.25 6.37 2.85 2.84 0.9 3.23 1.03 4.08 2.23 2.72 14.26 12.71 100.11 99.84 100.26 100.765 99.51 99.61 99.58 99.73 99.5 99.55 99.75

REE and trace element (ppm) 2.34 2.79 1.18 1.03 0.59 2.77 0.89 17.9 0.82 0.57 0.51 6.02 7.76 3.07 2.4 1.16 7.71 2.06 39.1 1.86 1.24 1.12 0.77 1.05 0.4 0.27 0.12 1.05 0.26 4.94 0.26 0.18 0.17 3.8 5.56 2.06 1.27 0.46 5.54 1.31 22 1.41 1.02 0.97 1.19 1.77 0.75 0.35 0.12 1.79 0.4 4.67 0.52 0.22 0.2 0.43 0.6 0.27 0.32 0.067 0.61 0.42 1.54 0.39 0.03 0.02 1.46 2.24 1.04 0.6 0.17 2.25 0.51 4.6 0.76 0.27 0.25 0.27 0.42 0.22 0.1 0.034 0.42 0.093 0.72 0.15 0.036 0.033 1.86 2.85 1.56 0.79 0.27 2.81 0.59 4.3 1.03 0.057 0.04 0.43 0.66 0.38 0.18 0.072 0.65 0.13 0.9 0.24 0.027 0.024 1.22 1.88 1.14 0.68 0.23 1.85 0.34 2.42 0.68 0.059 0.049 0.2 0.3 0.19 0.11 0.044 0.29 0.051 0.36 0.1 0.021 0.02 1.31 1.98 1.32 0.94 0.33 1.94 0.34 2.25 0.66 0.051 0.048 0.19 0.3 0.21 0.19 0.056 0.3 0.048 0.33 0.1 0.038 0.037 10.6 16.6 10.6 5.86 1.76 16.7 2.97 21.8 5.68 0.4 0.31 37.5 32.1 37.9 3.84 36.9 32.6 7.61 24.7 29.2 5.04 7.39 235 218 222 24.1 229 220 42.2 176 121 27.9 31.4 137 114 124 34.4 171 114 168 140 396 7920 2165 30 24.9 25.4 3.6 38.6 25 28.6 35.2 27.2 109 98.7 42.8 45.6 31.7 25.4 89.1 45.5 185 89.9 79.8 1684 2025 2.46 15 48.9 4.02 13.1 15.1 46.6 61.3 7.96 4.96 8.06 33.2 53.1 61.5 9.81 46.5 53.1 26.3 71.8 21.5 63.9 48.7 12.1 14.4 10.7 9.23 12.1 14.5 12.7 18.6 10.1 1.89 0.53 19.3 9.59 1.74 1.97 0.93 9.49 1.49 3.6 1.83 0.59 0.44 77.3 56.7 24 37.6 25.1 56.9 200 826 150 79.7 12.4 20.8 38.6 19 118 1.33 38.6 4.55 118 5.84 2.52 1.83 0.46 0.74 0.52 0.97 0.069 0.75 0.32 5.1 0.2 0.43 0.4 0.032 0.055 0.082 0.11 0.0091 0.055 0.03 0.32 0.019 0.082 0.085 0.63 0.078 0.079 0.18 0.087 0.078 0.037 0.38 0.076 0.071 0.057 66.7 104 37.6 37.4 15.1 104 16 299 27.5 14 3.74 0.73 1.25 0.62 3.21 0.065 1.25 0.18 3.15 0.24 0.037 0.02 1.49 1.36 1.88 13.3 1.41 1.36 1.64 4.07 1.28 1.17 1.15 0.52 0.57 0.24 0.29 0.1 0.56 0.13 1.58 0.11 0.098 0.09 0.17 0.19 0.18 0.11 0.025 0.19 0.047 0.46 0.068 0.077 0.035

(continued on next page) oscillatory and patchily zoned (e.g., Fig. 6-9, left three), as those of the age of the hornblendite. The younger cluster has a weighted mean age anorthosite. However, the remaining zircons are more or less irregular of 463 ± 4 Ma (n = 6, χ2 = 0.96), which, however, is determined by in shape (e.g., Fig. 6-9, right). Some of them partially preserve oscillatory the physically and chemically modified zircons and is therefore consid- or patchy zoning (Fig. 6-9; right top two), and some of them display ered geologically meaningless. stripes (e.g., Fig. 6-9, right bottom two), suggesting variable chemical modification. In the Concordia diagram (Fig. 9B), all the isotopic compo- 5.3.2. Diorites (tectonic blocks) in metavolcano-sedimentary mélanges sitions (Table S1) are concordant or near-concordant, but statistically (Zoolen) define two clusters. The older cluster has a weighted mean age of Zircons from diorite sample ZL10-1 are relatively euhedral, short- 512 ± 4 Ma (n = 7, χ2 = 1.85; Fig. 9B), corresponding to the rela- prismatic crystals and their CL images (e.g., Fig. 6-10) show broadly tively euhedral zircons. This age is slightly younger than, but broadly spaced and patchily zoned patterns, similar to gabbroic zircons similar to, the 519 ± 4 Ma age of the anorthosite (hosted in (e.g., Corfu et al., 2003). With one statistical exception (marked in hornblendite), and we interpret it to represent the minimum protolith gray), fifteen analyses are concordant and cluster at 520 ± 5 Ma Author's personal copy

80 P. Jian et al. / Earth-Science Reviews 133 (2014) 62–93

Table 2 (continued)

Locality Gurvan Sayhan Zoolen

Unit Hornblendite–anorthosite block Serpentinite mélange Diorite blocks

Sample no. ZL03-1 ZL03-2 ZL12-2 ZL12-1 ZL10-1 ZL10-2 ZL09-1 ZL09-2

Lithology Hornblendite Anorthosite Serpentinite Tonalite Diorite Tonalite Diorite Diabase (≥512 ± 4 Ma) (519 ± 4 Ma) (metaharzburgite–dunite) (dike; 496 ± 6 Ma) (520 ± 5 Ma) (dike; 449 ± 6 Ma) (511 ± 5 Ma) (dike)

Oxides (%)

SiO2 51.82 51.72 40.43 61.03 55.25 60.48 57.14 51.06

TiO2 0.56 0.24 b0.01 0.43 0.54 0.88 0.71 1.24

Al2O3 15.17 20.05 0.66 15.57 16.34 22.13 15.91 13.9 T Fe2O3 9.65 5.51 5.83 6.51 8.25 0.49 7.75 11.6

Fe2O3 FeO MnO 0.16 0.06 0.11 0.09 0.12 0.02 0.12 0.15 MgO 6.3 2.7 39.3 2.95 4.91 0.28 4.16 6.93 CaO 10.29 12.63 0.17 4.05 6.7 3.81 6.36 7.14

Na2O 3.7 3.87 0.08 5.87 4.27 8.78 4.9 4.84

K2O 0.3 0.2 0.01 0.23 0.83 1.53 0.34 0.18

P2O5 0.16 0.25 0.01 0.18 0.13 0.01 0.21 0.13

H2O

CO2 LOI 1.76 2.28 12.98 2.63 2.26 1.15 1.95 2.48 Total 99.87 99.51 99.58 99.54 99.6 99.56 99.55 99.65

REE and trace element (ppm) La 9.55 8.39 0.54 16.8 11.3 23.4 18.1 3.62 Ce 22.9 18.8 1.16 36.9 26.8 56.5 40 11 Pr 3.25 2.53 0.18 4.68 3.67 6.29 5.17 1.86 Nd 15.9 11.5 1.01 20.2 16.9 23.6 23.6 9.75 Sm 3.76 2.34 0.22 3.98 3.62 4.02 5.08 3.13 Eu 1.16 0.9 0.03 1.35 1.14 1.67 1.5 1.06 Gd 3.45 2.09 0.26 3.43 3.35 3.69 4.58 3.99 Tb 0.49 0.28 0.035 0.45 0.47 0.49 0.63 0.73 Dy 2.72 1.45 0.049 2.27 2.52 2.6 3.35 4.73 Ho 0.52 0.28 0.026 0.44 0.49 0.5 0.66 1.02 Er 1.45 0.79 0.055 1.19 1.35 1.41 1.83 2.94 Tm 0.22 0.12 0.022 0.18 0.2 0.21 0.27 0.45 Yb 1.3 0.73 0.05 1.12 1.26 1.27 1.63 2.76 Lu 0.19 0.11 0.039 0.17 0.19 0.19 0.25 0.42 Y 14 7.99 0.36 12.4 13.4 14.5 18.9 27.8 Sc 35.8 10.3 7.33 15.6 28.8 3.27 24.8 45.6 V 321 130 33.7 155 246 36.5 228 323 Cr 130 55.1 2194 10.1 36.5 48 41.7 101 Co 33 11.8 101 14.9 24.5 0.63 21.8 38.5 Ni 39.8 23.5 2057 8.43 19.7 1.93 20.5 52.3 Cu 174 36.5 7.84 96.5 7.29 5.11 84.2 115 Zn 62.7 20.6 50.9 45.3 50.4 5.56 69.8 73 Ga 17.9 21.6 0.53 16 16.3 12.1 18.9 12.5 Rb 2.92 2.13 1.5 4.55 7.95 10.1 2.96 1.42 Sr 726 1121 13.8 271 463 389 451 234 Zr 30 78.6 2.07 85.3 67.1 112 106 71.7 Nb 1.61 1.17 0.4 3.07 2.27 2.92 4.04 1.83 Ta 0.074 0.055 0.092 0.16 0.12 0.16 0.22 0.11 Cs 0.14 0.053 0.06 0.31 0.28 0.29 0.27 0.35 Ba 75.1 115 4.24 195 279 394 118 151 Hf 0.99 1.68 0.028 2.14 1.88 2.88 2.71 1.9 Pb 1.89 2.64 1.12 1.87 1.5 1.01 2.11 1.03 Th 0.38 0.45 0.098 1.76 1.48 2.95 2.42 0.28 U 0.22 0.4 0.04 0.85 0.43 2.03 1.05 0.16

(χ2 = 1.24; Fig. 9C), which we interpret as the time of diorite zoned (e.g., Fig. 6-12). With two statistical exceptions (marked in crystallization. gray), the crystallization age is determined as 496 ± 6 Ma (n = 8, Zircons separated from diorite sample ZL09-1 are uniformly large χ2 =0.41;Fig. 9E).Thisagerepresentsthetimeoftonalitedike (ca. 100 μm), euhedral short-prismatic crystals with oscillatory and emplacement. patchy zoning (e.g., Fig. 6-11). Their concordant ages are grouped at Most of the zircons from another tonalite dike (cutting a diorite 511 ± 5 Ma (n = 13, χ2 = 0.39; Fig. 9D) that is interpreted as the block) sample ZL10-2 crystallized at 449 ± 6 Ma (n = 8, χ2 = 1.46; time of diorite crystallization. Fig. 9F) that may represent dike emplacement. However, two zircons have younger ages of ca. 377 and ca. 388 Ma. Their CL patterns 5.3.3. Tonalite dikes (Zoolen) (e.g., Fig. 6-13, right bottom) are indistinguishable from those of Tonalite dike (cutting a serpentinite mélange) sample ZL12-1 the majority (e.g., Fig. 6-13; the remaining three). We are thus unable contains a single population of magmatic zircons that are oscillatory to propose a deﬁnite geological interpretation, but speculate that Author's personal copy

P. Jian et al. / Earth-Science Reviews 133 (2014) 62–93 81

these young zircons are related to a late thermal event. This dike is too 12

Nd(t) young to link with the orogeny in central CAOB and is precluded from − ε further discussion. Sr

86 5.3.4. Summary

Sr/ We have identiﬁed several early Cambrian tectonic blocks, including 87 a hornblendite (N512 ± 4 Ma; Table 1)withananorthositepatch (519 ± 4 Ma) and two diorites (520 ± 5 Ma; 511 ± 5 Ma) from the broadly volcano-sedimentary mélanges. The ophiolite has not yet been directly dated, but the 496 ± 6 Ma age of a tonalite dike cutting a serpentinite mélange constrains an upper age of ophiolite formation to the early Cambrian. Below we provide geochemical data to postulate a likely genetic link between the Gurvan Sayhan–

Sr Initial –

86 Zoolen ophiolite and the hornblendite diorite blocks (Section 6.3). Sr/ 87 6. Geochemical data

New geochemical (and Nd–Sr isotopic) data are given in Tables 2 and 3. Most of our samples, however, are variably evolved/fractionated plutonic rocks, which are altered or metamorphosed as indicated Nd

144 by their high LOI values (Loss on Ignition; up to ca. 8%; Table 2).

Nd/ The following report, which mainly concerns the geochemical 512916 ± 0.000008 0.705082 ± 0.000020 0.7048 4.3

143 constraints on magma sources, therefore addresses the initial Nd isotopic compositions (Table 3) that are insensitive to magmatic

Sr fractionation (e.g., McCulloch et al., 1980), as well as HFSE (high 86 ﬁeld-strength elements), REE (rare earth elements) and some tran- Rb/ 0.08 0.512707 ± 0.000009 0.712287 ± 0.000012 0.7119 4.3 0.4 0.512661 ± 0.000012 0.706544 ± 0.000022 0.7033 5.2 87 sitional elements (e.g., Cr, Ni), which in general remain virtually immobile during alteration and low-grade metamorphism (e.g., Pearce,

Nd 1983). 144

Sm/ 6.1. Dariv and Khantaishir (Northwest Mongolia) 15 0.13 0.512724 ± 0.000009 0.705450 ± 0.000015 0.7044 5.3 0.17 0.15 0.512890 ± 0.000010 0.704933 ± 0.000020 0.7037 6.6 147 6.1.1. Dariv

6.1.1.1. Microgabbro, isotropic gabbro and plagiogranite. One microgabbro

(sample MDRV09-1; 568 ± 5 Ma; SiO2 = 47.0 wt.%; Table 2)froma sheeted dike complex and two isotropic olivine gabbros (samples

MDRV07-1 and MDRV07-2; SiO2 = 47.0–47.2%) from a gabbro– pyroxenite suite are extremely poor in TiO2 (0.07–0.11%), but rich in MgO (8.8–16.7%), Cr (164–791 ppm) and Ni (73–262 ppm) and have FeOT/(FeOT +MgO)ratios(mafic index) of 0.41–0.81. They have extremely low REE concentrations (ca. 1–2 times chondrite), and their REE patterns are almost flat and smooth, without Eu anomalies to indi- Rb Sr Nd Sm ppm cate significant plagioclase accumulation or fractionation (Fig. 10A). The NMORB-normalized trace element patterns (spidergrams; Fig. 10B) display strong LILE (large-ion lithophile element) enrichment and highly depleted HFSE (ca. 0.1 times NMORB). These rocks also bear other typical arc signatures (e.g., Kelemen et al., 1993), i.e. the pronounced, consistently negative Nb–Ta–Zr–Hf–Ti anomalies (Fig. 10B). Overall,

512 ± 4 Ma 2.12 547 14the diagnostic 3.38 0.15 geochemical 0.01 features 0.51286 ± 0.00008 0.703892 ± of 0.000018 refractory 0.7038 maficmagmas 7.3 —high Ma ≥ Mg–Cr–Ni, coupled with extremely low Ti, REE and HFSE (e.g., Serri, 1981; Shervais, 2001), combined with the typical arc geochemical signatures, fingerprint a subduction-modified refractory mantle (wedge) source for these essentially non-cumulate gabbroic rocks. Indeed, ex-

cept for the low SiO2 contents, these rocks are very similar to boninites (typically, SiO2 N 53.6%) of intraoceanic forearcs (e.g., Hickey and Frey, 1982; Stern et al., 1991) in terms of their MgO, TiO2, Cr and Ni contents as well as their REE and trace element patterns. The dated plagogranite (sample MDRV09-3; 567 ± 4 Ma) has high ZL12-1 Tonalite 494 ± 6 2.52 235 20.3 3.99 0.12 0.03 0.512781 ± 0.000010 0.704746 ± 0.000020 0.7045 7.7 MDRV09-2M99D19 Diorite Granite Assumed at 570 Ma 487 ± 7 1.83 124 5.4 44.5 1.39 274 0.16 31.6 6.21 0.12 0.47 0.51179 ± 0.00007 0.711872 ± 0.000015 0.7086 ZL03-1 Hornblendite MHTS02-2MHTS06 GabbroMALT14 Andesite Gabbro Assumed at 570 Ma 2.89 518 ± 6 20.9 0.41SiO 0.09 40.4 1.312 0.13 (71.2%; 109 46.3Table 4.76 1.2 1.45 2 0.42), 0.18 moderate 0.21Al 2.53 0.032O3 (12.6%) 0.512909 ± 0.000009 0. 0.709406and ± 0.000010 extremely 0.6889 low 6.2 K2O (0.17%) contents. Such a chemical feature is characteristic of oceanic plagiogranites (Coleman and Peterman, 1975) or SSZ ocean ridge granites (ORGs) (Pearce et al., 1984b). This sample also displays a pronounced negative Eu anomaly (Fig. 10A) indicating signiﬁcant Sr isotopic compositions of selected samples.

– plagioclase fractionation. However, the plagiogranite is special in having Locality Sample no. Lithology Age Zoolen ZL09-1 Diorite 511 ± 5 2.82 345 22.5 4.82 0.13 0.02 0.512792 ± 0.000012 0.70407 ± 0.000015 0.7039 7.3 Dariv MDRV09-3 Plagiogranite 567 ± 4 8.08 177 3.33 0.8 0. Gurvan Sayhan ZL03-2 Anorthosite 519 ± 4 1.72 811 8.78 1.82 0.13 0.01 0.512805 ± 0.000009 0.704407 ± 0.000015 0.7044 7.9 Gobi Altai MALT18-1 Gabbro 523 ± 5 0.51 129 0.96 0.29 0.18 0.01 0.512962 ± 0.000009 0.702872 ± 0.000010 0.7028 7.4 Khantaishir M98K12 Plagiogranite 565 ± 7 1.89 35.8 1.49 0.42

Table 3 Nd MgO (1.1%) content somewhat higher than the classic SSZ ORGs (i.e. in Author's personal copy

82 P. Jian et al. / Earth-Science Reviews 133 (2014) 62–93 AB

G H

I J

Fig. 10. Chondrite-normalized REE patterns (left) and NMORB-normalized multi-element variation diagrams (right). A and B—samples from Dariv; C and D—samples from Khantaishir; EandF—An amphibolite (MHTS07-1) from a thrust/shear zone between the Khantaishir Ophiolite and the Dzabkhan–Baydrag microcontinent; G and H—Samples from Gobi Altai; I—samplesfromGurvanSayhanandZoolen;J— a diabase (ZL09-2) from Zoolen. Chondrite and N-MORB values after SunandMcDonough(1989).

the Tuscany, Smarville and Troodos ophiolites; Pearce et al., 1984b) that crustal recycling by oceanic subduction, the plagiogranite is deplet- contain less than 0.6% MgO. Another striking feature is that, except ed in all of the elements either in LILE (e.g., K, Rb and Ba) or in HFSE for Th, which Pearce (2008) highlights as a sensitive indicator of (e.g., Nb, Ta, Zr, Sm, Y and Yb) relative to the SSZ ORG average values Author's personal copy

P. Jian et al. / Earth-Science Reviews 133 (2014) 62–93 83

Fig. 11. ORG-normalized multi-element variation diagrams. A—Dariv and Khantaishir plagiogranites; B—Gurvan Sayhan and Zoolen samples (anorthosite, hornblendite, diorite and tonalite). ORG values after Pearce et al. (1984b).

(Pearce et al., 1984b; Fig. 11A). The subduction signatures are An isotropic gabbro (sample MHTS02-2, SiO2 = 48.3%), a dacite reflected by the enrichments in LREE (Fig. 10A) and LILE and nega- (sample MHTS04; SiO2 = 63.1%; Table 2) and the dated plagiogranite tive Nb–Ta anomalies (Fig. 11A). These geochemical features again (sample M98K12, 566 ± 7 Ma; SiO2 = 77.5%), exhibit similar REE suggest a subduction-modified refractory mantle (wedge) source patterns with positive Eu anomalies (Fig. 10C) suggesting plagioclase for the plagiogranite, as for the microgabbro and isotropic gabbros. accumulation (Fig. 10C). With increasing SiO2 contents the REE concen- Melting of refractory mantle wedge is believed to match an early trations show an increase, which is predicted as a consequence of phase of SSZ ophiolite formation and to signify a transition from the fractionation of a basaltic magma (e.g., Brophy, 2008). Therefore, the subduction initiation to the establishment of an incipient volcanic arc plagiogranite possibly crystallized from a residual liquid, left after pro- complex (Shervais, 2001). gressive fractionation of a basaltic magma, as proposed by Coleman and Peterman (1975). 6.1.2. Diorite and leucogranite dikes 6.1.3.2. An amphibolite on the ophiolite/microcontinent boundary. This Cross-cutting diorite dike (sample MDRV09-2; SiO = 54.2%; 2 dated amphibolite (sample MHTS07-1, 514 ± 8 Ma; Table 2)hasa Table 2) in the sheeted dike complex has high Mg (MgO = 6.6%), basaltic composition. The REE pattern is flat with a negative Eu Cr (391 ppm) and Ni (71 ppm) contents and is geochemically similar anomaly (Fig. 10E). The rock is enriched in LILE, and depleted in to a sanukitoid high-Mg andesite in western Japan (e.g., Tatsumi and Nb–Ta (Fig. 10F). This might be misinterpreted as an arc rock solely Ishizaka, 1982; Shimoda et al., 1998). The diorite displays LREE- in terms of geochemical criteria, but fortunately, the presence of (Fig. 10A) and LILE-enriched (Fig. 10B) patterns with pronounced abundant detrital zircons suggests a sedimentary protolith (see negative Nb–Ta and Ti anomalies. It has a positive ε value of Nd(t) Section 5.1.2.2). 4.3 (Table 2) suggesting a depleted mantle source. Leucogranite dike (sample M98D19; 487 ± 7 Ma; SiO = 72.7%; 2 6.1.4. Nd–Sr isotopic compositions of selected samples Table 2)inthegabbro–pyroxenite complex is K- (K O=3.1%) 2 An andesite lava (sample MHTS06), an isotropic gabbro (sample and LREE-enriched (Fig. 10A). The Nd–Sr isotopic compositions MHTS02-2), together with two dated plagiogranites (samples (ε = −11.7, initial 87Sr/86Sr ratio = 0.7086; Table 3) suggest Nd(t) M98K12 and MDRV09-3), were measured for Nd–Sr isotopic com- a predominantly crustal origin. This interpretation is supported positions (Table 3). These rocks have limited dispersion in initial by the presence of Neoproterozoic zircon xenocrysts in the rock Nd isotope compositions (ε =5.2–6.6), independent of lithol- (see Section 5.1.1.3). Nd(t) ogy and structural position. The relatively uniform and positive

εNd(t) values provide robust evidence that the andesite (6.2; 6.1.3. Khantaishir Table 3), isotropic gabbro (5.2) and plagiogranites (6.6 and 5.3) originated from essentially the same depleted mantle-derived 6.1.3.1. Lavas, an isotropic gabbro and a plagiogranite. The lavas from the melt that generated the Dariv–Khantaishir ophiolitic crust. This serpentinite mélange show a large compositional range (andesitic ba- conclusion is partly supported by the initial 87Sr/86Sr ratios, which, salt–andesite–dacite) (Table 2). The andesitic basalt (sample MHTS01; with one exception, are within a narrow range of 0.7033–0.7044

SiO2 = 54.2%; TiO2 = 0.37%; Table 2) displays a near-ﬂat REE pattern (Table 3).Theandesite(sampleMHTS06)hasanexceptionalandun- and has lower REE and HFSE concentrations (Fig. 10C and D) than reasonably low initial 87Sr/86Sr ratio of 0.6889. This rock has a high Rb NMORB values, broadly consistent with an island arc tholeiite (IAT) content (40.36 ppm) that may be one of the likely causes of this

(e.g., Hawkins, 2003). The andesite (sample MHTS06, SiO2 =57.6%; problem. TiO2 =0.59%;Table 2)alsoshowsanear-flat REE pattern but has higher Our newly obtained εNd(t) values (5.2–6.6), combined with REE and HFSE concentration levels than the basaltic andesite (Fig. 10C those previously reported values for two Dariv plagiogranites and D). The observed trend, i.e. increasing REE and HFSE with SiO2,is (5.4–5.8; Khain et al., 1995; Kozakov et al., 2002)andone readily explained by magma differentiation (e.g., Brophy, 2008). These Khantaishir plagiogranite (8.4; Kozakov et al., 2002), are low two lavas are enriched in LILE and exhibit pronounced negative Nb–Ta when compared with the 570 Ma NMORB value (8.8; calculated anomalies (Fig. 10D), reflecting subduction signatures. On the other after Nelson and DePaolo, 1984). Correspondingly, the initial 87Sr/86Sr hand, the two lavas have relatively high Cr (114–124 ppm), Ni (43–46 ratios (0.7033–0.7044; Table 3) are elevated relative to NMORB ppm) concentrations and are otherwise similar to boninites from (modern average = 0.7025; McCulloch and Perfit, 1981). Lower intraoceanic forearcs (Cr, mostly N 200 ppm; Ni N 70 ppm; Hickey and εNd(t) values than contemporaneous NMORBs have been obtained Frey, 1982; Stern et al., 1991). Accordingly, we conclude that the two for many SSZ-type-ophiolite magmatic rocks—the Cretaceous lavas represent variably evolved IAT-boninitic magmas that bear (NMORB value ≈ 10) Semail (7.6–8.2; McCulloch et al., 1980) diagnostic SSZ geochemical signatures (e.g., Shervais, 2001). and Troodos (7.9–0.8; McCulloch and Cameron, 1983)arethe Author's personal copy

84 P. Jian et al. / Earth-Science Reviews 133 (2014) 62–93 best-known examples. With broad consensus, this is accounted for effect (e.g., Pearce, 2003). The elevated initial 87Sr/86Sr ratios are by the interaction between a subducting slab-derived LREE- alsocommoninSSZ-typeophiolites but can be interpreted differ- enriched component and an overlying mantle (e.g., McCulloch ently—either indicate oceanic subduction (e.g., Shervais, 2001)or and Cameron, 1983), and is consequently ascribed to a subduction reﬂect sub-ocean ﬂoor alteration (e.g., McCulloch et al., 1980) Author's personal copy

P. Jian et al. / Earth-Science Reviews 133 (2014) 62–93 85

depending on the specific situation. However, decreasing εNd (t) (Yb = 1.3 ppm) and HFSE (Fig. 10I; Fig. 11B), with a negative Eu anom- values with increasing initial 87Sr/86Sr ratios relative to NMORBs, aly (Fig. 10I) suggesting plagioclase removal during magma evolution. It as in the present case, have been documented by Shervais (2001) is high in Sr (726 ppm), low in Y (14 ppm), resulting in a high Sr/Y ratio as a general isotopic feature of subduction zone magmatic rocks (52) and thus reflecting an adakitic trait (Sr/Y N 20, Y ≤ 18 ppm; in SSZ-type ophiolites. Martin, 1999). However, the Lan/Ybn (5.3) ratio (adakite = 14.2) and K2O (0.3%) content (adakite = 1.72%) are low; and conversely, the 6.2. Gobi Altai (South Mongolia) MgO (6.3%; Table 2) (adakite = 2.2%), Cr (130 ppm) (adakite = 36 ppm) and Ni (40 ppm) (adakite = 24 ppm) contents are high, The two dated gabbros (Table 2) are cumulates, as indicated by when compared with an average adakite (Martin, 1999). Instead, their pronounced positive Eu anomalies (Fig. 10G). Layered gabbro these compositional features, combined with the extremely low HREE N sample MALT18-1 (523 ± 5 Ma) has an ε value of 7.4 (Table 3), and HFSE concentrations, are more akin to boninites (SiO2 53.6%; Nd (t) N – b – slightly lower than, but close to the Early Cambrian NMORB (8.9; MgO 7.76%; K2O=0.3 1.47%, mostly 0.51%; TiO2 =0.1 0.5%, N N N after Nelson and DePaolo, 1984). The initial 87Sr/86Sr ratio of this Cr 74 ppm, mostly 200 ppm; Ni 70 ppm; Hickey and Frey, 1982; sample is 0.7028 (Table 3) and is slightly more elevated than the Stern et al., 1991). We therefore conclude that the hornblendite origi- – fi average NMORB value (0.7025; McCulloch and Perfit, 1981). nated from a hybrid adakite boninite melt. By de nition, adakite is de- Leucogabbro sample MALT14 (518 ± 6 Ma) has an even lower rived by slab melting where garnet is stable (Defant and Drummond, 87 86 1990; Martin, 1999; Martin et al., 2005), and boninite mostly forms by εNd(t) value of 4.3, and correspondingly, a higher Sr/ Sr ratio of 0.7048. The Nd–Sr data, although scarce, point to a likely trend of melting of a refractory mantle wedge above a subduction zone (e.g., 87 86 Crawford et al., 1981; Hickey and Frey, 1982; Stern et al., 1991). decreasing εNd(t) value with increasing initial Sr/ Sr ratio, as in the Dariv–Khantaishir ophiolite. The anorthosite patch (sample ZL03-2; 519 ± 4 Ma; SiO2 =51.7%; Table 2) in hornblendite is a plagioclase cumulate with a high Al2O3 con- Diabase dike sample MALT18-2 (SiO2 =47.8%; Table 2)is strongly LREE-enriched (Fig. 10G) and is inconsistent in geochem- tent (20.1%). This rock shows LREE- (Fig. 10I) and LILE-enriched istry with an ophiolite. This is most likely a post-ophiolite dike. (Fig. 11B) patterns similar to those of the host hornblendite and displays There are no reported geochemical data for peridotites and basalts, a pronounced positive Eu anomaly (Fig. 10I), complementary to the and the genetic type of the Gobi Altai ophiolite is currently difficult negative Eu anomaly of hornblendite. The anorthosite is therefore co- to constrain. magmatic with the hornblendite.

6.3. Gurvan Sayhan and Zoolen (Trans-Altai, South Mongolia) 6.3.3. Diorites (tectonic blocks) and a diabase dike

Diorite samples ZL10-1 (520 ± 5 Ma; SiO2 = 55.3%; Al2O3 =15.6%; 6.3.1. Metaperidotites (serpentinites) Table 2) and ZL09-1 (511 ± 5 Ma; SiO2 =57.1%;Al2O3 = 15.9%) have The three serpentinites (samples ZL01-2, ZL02-1 and ZL12-1; Table 2) MgO (4.9%; 4.2%), Cr (37 ppm; 42 ppm) and Ni (20 ppm; 21 ppm) con- – – consistently have high MgO (ca. 35 40%), Cr (2165 7920 ppm) and Ni tents, which are intermediate between adakite and boninite like the – – (1684 2057 ppm), coupled with low Al2O3 (0.66 0.89%) and extremely hornblendite described above. They also have high Sr (463 and b – low TiO2 ( 0.01%), and are metaharzburgite dunite. The chemically 451 ppm), low Y (13.4 and 18.9 ppm), resulting in high Sr/Y ratios – documented harzburgites dunites, together with previously reported (35 and 24), like an adakite (Martin, 1999). The two diorites display lherzolites in the area (Zonenshain and Kuzmin, 1978), constitute a moderate LREE-enriched patterns (Lan/Ybn = 6.5; 7.9) without signiﬁ- peridotite association similar to that in some ancient and modern forearcs cant Eu anomalies, indicating neither plagioclase fractionation nor accu- (e.g., Pearce et al., 2000). Therefore in the following discussion we consid- mulation (Fig. 10I). These features are thus consistent with hybrid – er the Gurvan Sayhan Zoolen ophiolite as a SSZ-type, i.e. a forearc adakite–boninite melts. Moreover, compared with ORG, these rocks basement. are enriched in LILE and severely depleted in HFSE (Fig. 11B), indicating a refractory mantle source. Similar trace element patterns have 6.3.2. Hornblendite–anorthosite (a tectonic block) been documented for the Dariv–Khantaishir plagiogranites (Fig. 11A),

The hornblendite (sample ZL03-1; ≥512 ± 4 Ma; SiO2 = 51.8%; which are interpreted to have originated from a mantle wedge beneath Table 2) has a basaltic andesite composition and is low in K (K2O= an incipient forearc. 0.3%) and Ti (TiO2 = 0.56%). This rock shows moderate LREE (Lan/ A diabase dike (sample ZL09-1, SiO2 = 51.6%; Table 2) cutting Ybn = 5.3) and LILE enrichments over extremely depleted HREE diorite has a LREE-depleted NMORB-like REE pattern (Fig. 10I). It is

Fig. 12. Cross-sections illustrate the timing of orogeny in the central CAOB in the latest Neoproterozoic-Cambrian. (A) Primary setting (ca. 655–636 Ma)—an open ocean with drifting microcontinents. The age of the Bayankhongor MORB-type ophiolite (ca. 655–636 Ma; Fig. 1A; Jian et al., 2010a) is assumed to represent the opening of this ocean—the late Neoproterozoic Paleo-Asian Ocean (e.g., Khain et al., 2003; Windley et al., 2007; Wilhem et al., 2012). Shown on the left is future South Mongolia, on the right future Northwest Mongolia, which is separated by the future Main Mongolian Lineament (MML, Tomurtogoo, 1997; Xiao et al., 2004; Windley et al., 2007). (B) Phase 1 (ca. 573 to N ca. 540 Ma)—Subduction initiation and arc formation in Northwest Mongolia. Subduction initiation is suggested by the generation of the Dariv–Khantaishir SSZ-type ophiolite (ca. 573–560 Ma; this work), which was closely followed by the formation of the Lake arc (ca. 551–540 Ma; loosely conﬁned according to Kröner et al., 2001; Kovalenko et al., 2004; Kovach et al., 2011; Yarmolyuk et al., 2011; Rudnev et al., 2012). (C) Phase 2 (ca. 535–524 Ma)—Syn-collisonal magmatism in the Lake arc. There are two likely explanations for the generation of the ca. 551–524 Ma high-Al TTG-like suite (Rudnev et al., 2009, 2012; Kovach et al., 2011; Yarmolyuk et al., 2011) that required a heat source most likely from depth. Interpretative selection 1: by ridge-trench collision, to create a slab window that allowed upwelling of asthenosphere and resultant melting of the lower arc crust. Interpretative selection 2: asthenosphere upwelling due to slab breakoff, followed by forearc– microcontinent collision. (D) Phase 3 (ca. 519–482 Ma)—A continuum of slab delamination, overthrusting, crustal thickening and surface uplift. The published age data are marked by a low-Al TTG-like suite (ca. 519–494 Ma) and alkaline melts (alkaline granite and peridotite–pyroxenite–granodiorite) (ca. 511 Ma) in the Lake arc (Rudnev et al., 2009; Kovach et al., 2011; Yarmolyuk et al., 2011; Rudnev et al., 2012); and in the Dazbkhan–Baydrag microcontinent for high-grade metamorphism (ca. 510 Ma; Kozakov et al., 2002) and broadly granitoid magmatism (ca. 515–482 Ma; Dijkstra et al., 2006; Kröner et al. unpublished). Thrusting of the Dariv–Khantaishir ophiolite and the Lake arc over the Dazbkhan–Baydrag microcontinent at ca. 514 Ma is interpreted from the present work. For detailed explanations of magmatic and magmatic expressions of tectonic events see the text. (E) Phase 4 (ca. 523–511 Ma)—Initiation of new subduction zone(s) in South Mongolia. Left: the generation of the Gurvan Sayhan–Zoolen forearc (ca. 520–511 Ma; this work) with an ophiolite basement (Nca. 494 Ma). Formation of the Gobi Altai ophiolite (ca. 523–518 Ma; this work) has two possibilities: on the right, we suggest a new subduction zone before the Gobi Altai microcontinent; in the middle, we speculate spreading of a back-arc basin behind the Gurvan Sayhan–Zoolen arc. (F) Phase 5: Continuing orogeny with local surface uplift. Left: migration of subduction southwards. Right: obduction of the Gobi Altai ophiolite onto the Gobi Altai microcontinent. Also marked are a tonalite dike (ca. 494 Ma; this work) cutting the Gurvan Sayhan–Zoolen ophiolite, and broadly granitoid magmatism (ca. 518–498 Ma; Hanžl and Aichler, 2007; Hrdličková et al., 2008; Demoux et al., 2009b; Kröner et al., 2010; Hrdličková et al., 2010) in the Gobi Altai microcontinent. Left: development of accretionary complex in the South Gobi zone-Northern orogen of Chinese Inner Mongolia in response to southward migration of subduction (ca. 500 Ma; Jian et al., 2008). Author's personal copy 86

Table 4 Summary of age and geochemical data that interpret the latest Neoproterozoic to Cambrian orogenesis in central CAOB.

Area/unit Northwest Mongolia

Lake zone Ophiolite

Volcano Conglomerate Granitoids Peridotite– Dariv Khantaishir -sedimentary pyroxenite– sequence anorthosite–

gabbronorite 62 (2014) 133 Reviews Earth-Science / al. et Jian P.

Association Basalt– Silicious– Quartz Granite Diorite– Diorite– Granite– Diorite– Gabbronorite Plagiogranite– Agranitedike Plagiogranite An amphibolite and andesite– terrigenous porphyry (clast) tonalite– tonalite– monzodiorite granodiorite– microgabbro (undeformed) in a thrust/shear dated dacite– sediments (clast) trondhjemite trondhjemite (alkaline) granite– cutting zone on the rock rhyolite (predominantly (predominantly leucogranite ophiolitic ophiolite/ high Al TTG- low-Al TTG- gabbro microcontinent like) like) boundary Zircon age 540 ± 1 Ma 492 ± 1 Ma The oldest, 519 ± 8– 511 ± 2 Ma 465 ± 11– 511±7Ma 568 ± 5– 487±7Ma 573 ± 8– 514 ± 8 Ma (207 Pb/ (207 Pb/ 551 ± 13 Ma; 494 ± (U–Pb TIMS; 456 (SHRIMP) 560 ± 8 Ma (SHRIMP) 565 ± (metamorphic age) 206 Pb; 206 Pb; mostly, 535 ± 10 Ma upper ±4Ma (SHRIMP) 5Ma (SHRIMP) evaporation) evaporation) 6– (SHRIMP) intercept) (SHRIMP) (SHRIMP) 524 ± 10 Ma (SHRIMP) Ar–Ar age 546 ± 3 Ma (amphibole) (from an andesite)

εNd(t) 9.9–7.3 7.4–5.6 9.0–7.4 7.9–6.5 7.2–6.0 6.9–3.1 5.3 5.4–5.8 −11.7 6.6 8.4 – 93 Reference Kovalenko et al. (2004), Kröner et al. (2001) Rudnev et al. (2009, 2012), This work Khain et al. This work Kozakov et This work Yarmolyuk et al. (2011), Yarmolyuk et al. (2011), (1995), al. (2002) Kovach et al. (2011) Kovach et al. (2011) Kozakov et al. (2002) Author's personal copy

Table 4 TableSummary 4 (continued of age and) geochemical data that interpret the latest Neoproterozoic to Cambrian orogenesis in central CAOB.

Area/unit Northwest Mongolia South Mongolia

Dzabkhan–Baydrag block Gobi Altai zone Trans-Altai zone

near Dariv Variably foliated Ophiolite Gurvan Sayhan and Zoolen ophiolitic mélanges granitoids

Association Garnet– An undeformed quartz Granitoid (foliated) Granite (undeformed) Diorite–granodiorite– Gabbro Hornblendite– Tonalite Tonalite (dike Intermidiate-felsic Siliceous and orthopyroxene diorite from a metavolcano- granite anorthosite and (dike cutting cutting a volcanic rocks (metrix) siltstones dated gneiss (i. sedimentary unit overlying a diorites (tectonic serpentinite) diorite block) and 62 (2014) 133 Reviews Earth-Science / al. et Jian P. rock e. a felsic crystalline continental blocks; high Sr/Y, volcaniclastic granulite) basement Cr–Ni) rocks (metrix) Zircon age 510 ± 4 Ma 515 ± 8 Ma (SHRIMP) 507 ± 4 Ma 498 ± 1 Ma (zircon 542±4Ma;mostly, 523 ± 5 Ma, 520 ± 5– 494 ± 6 Ma 449±6Ma 481 ± 10–417 ± 2 Ma Assumed at (TIMS U–Pb; evaporation); 518 ± 5–498 ± 3 Ma 518 ± 6 Ma 511 ± 5 Ma (SHRIMP) (SHRIMP) (SHRIMP) 420 Ma lower 482 ± 7 Ma (SHRIMP) (SHRIMP and ICP-MS) (SHRIMP) (SHRIMP) intercept) Ar–Ar age εNd(t) (−10.5 (−1.5 to −0.2) 7.4–4.3 7.9–7.3 7.7 7.3 (for a 9–67.4–6.3 to hornblendite) −8.2) Reference Kozakov et al. Dijkstra et al. (2006) Kröner et al. Kozakov Kröner et al. Hanžl and Aichler (2007), This work Helo et al. (2006),Kröner Helo et al. (2002) (unpublished); for et al. (unpublished); for Demoux et al. (2009b), et al. (unpublished), for (2006);see analytical data, see (2002) analytical data see Table Kroner et al. (2010), analytical data see Table also Table S2, sample no. S2,samplenos.99/D20 Hrdličková et al. (2008, S2, sample no. MO371) Yarmolyuk M99/D20 and M06/473 2010) et al. (2007) – 93 87 Author's personal copy

88 P. Jian et al. / Earth-Science Reviews 133 (2014) 62–93 enriched in LILE, and shows negative Nb–Ta anomalies with respect microcontinent collision zone in Northwest Mongolia (Fig. 1A), and to Th and La (Fig. 10J). This is typically a subduction zone signature the remaining two phases (Fig. 12E and F) for the orogenic events (e.g., Kelemen et al., 1993) and suggests an arc setting for emplace- in South Mongolia and geologically correlative areas in China. ment of the dike. 7.1. Phase 1 (ca. 573 to N ca. 540 Ma): subduction initiation and arc 6.3.4. A tonalite dike cutting ophiolite formation in Northwest Mongolia Tonalite sample ZL12-1 (496 ± 6 Ma; SiO2 =61.0%;Table 2)isalow K O volcanic arc granititoid rock (Y + Yb = 13.2 ppm; Rb = 1.4 ppm), 2 Our geochemical data for lavas and a microgabbro (see Section 6.1), judging from a plot of Rb versus Y + Yb (Pearce et al., 1984b;not together with previously reported geochemical data for mafic dikes of a shown). It is strongly enriched in LREE (La /Yb = 11) and LILE n n sheeted dike complex (Dijkstra et al., 2006), suggest that the crust of the over extremely depleted HREE and HFSE (Figs. 10I; 11B). Unlike Dariv–Khantaishir ophiolite was generated from a transitional IAT- the hybrid adakite–boninite magmatic rocks (i.e. hornblendite boninite magma. With analogous interpretations of IATs and boninites and diorite) from the volcano-sedimentary mélanges, this tonalite in the IBM arc in the W-Pacific(e.g.,Stern et al., 1991; Hawkins, 2003), dike cutting the ophiolite has low Sr (271 ppm), Cr (10 ppm) and Ni such a magma predicates a forearc setting for the Dariv–Khantaishir (8 ppm) concentrations. We suggest an arc tholeiitic precursor to ac- ophiolite (Fig. 12B). Moreover, the plagiogranites in the ophiolite have count for its major geochemical features (i.e. low K, HREE, HFSE, Cr, trace element patterns (e.g., Fig. 11B) and Nd isotopic compositions and Ni, etc.). (εNd(t) =5.3–8.4; Table 4; Khain et al., 1995; Kozakov et al., 2002)consistent with generation from a refractory mantle wedge, i.e. beneath an 6.3.5. Nd–Sr isotopic compositions incipient forearc (Shervais, 2001). The time of subduction initiation/ forearc spreading (Stern and Bloomer, 1992) in Northwest Mongolia is 6.3.5.1. Hornblendite–anorthosite–diorite (tectonic blocks). The therefore constrained to the period 573 ± 8 to 560 ± 8 Ma, i.e. the hornblendite, anorthosite and a diorite have limited dispersion in time of formation of the Dariv–Khantaishir SSZ-type ophiolite (Table 4). both initial Nd and Sr isotope compositions (Table 3)withε Nd(t) The Lake zone west of the Dariv–Khantaishir ophiolite (Fig. 1A) is ranging from 7.3 to 7.9 and initial 87Sr/86Sr ratios from 0.7038 to dominated by a BADR suite that represents an intra-oceanic arc 0.7044. Like the magmatic rocks of the Dariv–Khantaishir ophiolite, (Kovalenko et al., 2004; Kovach et al., 2011; Yarmolyuk et al., 2011). they show decreased ε and elevated initial 87Sr/86Sr ratios relative Nd(t) SHRIMP dating of zircons from granitoids in the Lake zone (Rudnev to contemporaneous NMORB (ε = 8.9; Nelson and NMORB (520 Ma) et al., 2009; Kovach et al., 2011; Yarmolyuk et al., 2011; Rudnev et al., DePaolo, 1984).Theserocksarethuscomparabletoaclassicintra- 2012) gave an oldest age of 551 ± 13 Ma interpreted by Rudnev et al. oceanic forearc in isotopic compositions, as in the Aleutians (ε = Nd (t) (2012) as the approximate formation age of the Lake arc. If this interpre- 6–9; initial 87Sr/86Sr, b0.7032 to 0.7045) (McCulloch and Perfit, 1981; tation is correct, then the formation time of the arc can be bracketed Drach et al., 1986). within the ca. 551–540 Ma range, taking into account the lower limit of ca. 540–546 Ma (Table 4; Kröner et al., 2001; Kovalenko et al., 6.3.5.2. The tonalite dike cutting ophiolite. ThetonalitedikesampleZL12-1 2004; Kovach et al., 2011; Yarmolyuk et al., 2011). It is thus most likely (Table 3) has an ε value of 7.7 and an initial 87Sr/86Sr of 0.7045. The Nd(t) that formation of the Lake arc (ca. 551–540 Ma) was immediately after rock has a juvenile nature, and this does not conflict with an origin the subduction initiation (ca. 573–560 Ma) (Fig. 12B). through melting in a newly erupted arc tholeiite.

7. Discussion: timing of orogeny in the central CAOB 7.2. Phase 2 (ca. 535–524 Ma): likely syn-collision granitoid magmatism in the Lake arc The zircon age and geochemical data reported above provide crucial constraints on the early development of oceanic subduction zones in Rudnev et al. (2009, 2012), Kovach et al. (2011) and Yarmolyuk et al. Northwest Mongolia north of the MML (Fig. 1A),aswellasonthe (2011) identified a predominantly high Al TTG (tonalite–trondhjemite– southern side in the Gobi Altai and Trans-Altai zones of South granodiorite; Jahn et al., 1988)-like suite—the diorite–tonalite– Mongolia. We can now establish the timing of orogeny in the cen- trondhjemite granitoid suite (ca. 535–524 Ma; Table 4)intheLake tral CAOB for five major evolutionary phases (Fig. 12) by incorpo- zone. This episode of magmatism seemingly did not affect the rating, from the literature, zircon ages and Nd isotopic data of Dzabkhan–Baydrag microcontinent (Fig. 12C). The high-Al TTG-like magmatic and metamorphic rocks in west Mongolia (Table 4)and granitoids have extremely high Sr/Y (107–287), coupled with high adjacent Chinese areas (Fig. 1A). The primary setting was an ocean- Lan/Ybn (mostly 8.1–19.0; Rudnev et al., 2009; Kovach et al., 2011; ic archipelago, namely the Paleo-Asian Ocean (e.g., Khain et al., Rudnev et al., 2012) ratios, and thus are broadly consistent geochemi- 2003; Windley et al., 2007; Wilhem et al., 2012) in the late cally with a plutonic adakite (e.g., Martin, 1999). Previous investigators Neoproterozoic (Jian et al., 2010a) with several drifting (Rudnev et al., 2009; Kovach et al., 2011; Yarmolyuk et al., 2011; Rudnev microcontinents—like the mid-Neoproterozoic (Zhao et al., 2006; et al., 2012) interpreted this suite as the plutonic part of the Lake arc. Yarmolyuk et al., 2008a; Levashova et al., 2010; Kozakov et al., However, these granitoids (ca. 535–524 Ma) intruded into the Lake vol- 2012)toArchean(Kotov et al., 1995; Kozakov et al., 2001, 2007; canic arc, and their emplacement was well after subduction initiation Demoux et al., 2009a,b)Dzabkhan–Baydrag microcontinent in (ca. 573–560) as well as volcanic arc formation (ca. 551–540 Ma, as future Northwest Mongolia, and the earliest Neoproterozoic to loosely defined above). This contradicts a corollary that all arcs form late Mesoproterozoic Gobi Altai microcontinent (Demoux et al., in a short time span (e.g., Shervais, 2001; Stern, 2002, 2004), and so 2009b; Kröner et al., 2010) in future South Mongolia (Fig. 12A). we propose the following two alternative interpretations. A late Neoproterozoic ocean, relevant to our discussion and Rudnev et al. (2009, 2012) and Kovach et al. (2011) proposed that already proposed by Khain et al. (2003),isindicatedbythepres- the high-Al TTG-like suite (εNd (t) =9.0–7.4; Table 4) likely originated ence of a ca. 655–636Ma,MORB-typeophiolite(Jian et al., 2010a) in the lower crust by partial melting of a mafic juvenile source at pres- in central Mongolia, immediately north of the Dzabkhan–Baydrag sures higher than 10 kbar. This requires a deep heat source to melt microcontinent (Fig. 1A). Note also the possible age distinction the lower arc crust. According to Kröner et al. (2010) there is an early between the Dzabkhan–Baydrag and Gobi Altai microcontinents. Cambrian unconformity in the Lake zone. Likely, the emplacement of As discussed below, the first three evolutionary phases (Fig. 12B, the high-Al TTG-like suite accompanied the relevant surface uplift C and D) account for the orogenic development of an arc– (Fig. 12C; top and bottom). Author's personal copy

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As shown in Fig. 12C (top), with reference to a Shervais (2001) (N494±6Ma;Table 4;thiswork;Fig. 12E, left). The polarity of this sub- model, as the forearc developed, a spreading ridge was supposedly duction zone is suggested by the fact that a huge accretionary complex approaching the trench and subsequently collided with it. This allowed occurs in the present south (for explanation see below). The dated rocks a slab window to open; subsequently, upwelling of asthenosphere from (a hornblendite, an anorthosite and two diorites) have hybrid adakite– the window may have heated up the lower arc crust and yielded the boninite affinities (see Section 6.3) and probably represent an early high-Al TTG-like magmas. These processes resulted in the cessation of phase of forearc development (Shervais, 2001). Thus, the subduction subduction, generation of the high-Al TTG-like suite and accompanied initiation theoretically occurred slightly earlier, and a future SHRIMP surface uplift. However, currently, there is no evidence for a ridge- study of the Gurvan Sayhan–Zoolen ophiolite is required to constrain trench collision in the Lake zone. the exact timing of subduction initiation. Initiation of new subduction Alternatively, a microcontinent was approaching and consequent zones, often with reversed polarities relative to the old failed ones in re- forearc–microcontinent collision (Fig. 12C; bottom) led to slab breakoff sponse to arc–microcontinent/oceanic plateau collisions is one of the (Davies and von Blanckenburg, 1995; Shervais, 2001), upwelling of as- most important tectonic processes in the modern SW-Pacific(Hall, thenosphere and consequent generation of the high-Al TTG-like suite, 2002; Stern, 2004), in particular, to the east of Australia at Papua, New followed by surface uplift. Accordingly, we link the high-Al TTG-like Caledonia and Northland (Whattam et al., 2008; Whattam, 2009). suite to arc–microcontinent collision. The lack of coeval magmatism in Meanwhile or slightly earlier, the Gobi Altai ophiolite (523 ± the colliding Dzabkhan–Baydrag microcontinent is not necessarily 5–518 ± 6 Ma; Table 4; this work) was formed, the original tectonic surprising. A general absence of syn-collisional magmatism in a lower setting of which is difficult to ascertain because of a major lack of (colliding) plate has long been noted by Pearce et al. (1990). geochemical data for peridotites and lavas. We propose two possibilities. First, this ophiolite likely formed in front of the Gobi Altai microcontinent 7.3. Phase 3 (ca. 519–482 Ma): a continuum of slab delamination, by initiation of a new south-dipping subduction zone (Fig. 12E, right) fol- overthrusting, crustal thickening and surface uplift lowing arc–microcontinent collision (ca. 535–524 Ma) in Northwest Mongolia, as in the Trans-Altai zone. Second, it may represent the resid- We determined the metamorphic age of 514 ± 8 Ma (Table 4)ina ual of a back-arc basin behind the Gurvan Sayhan–Zoolen arc (Fig. 12E; shear zone on the boundary between a Khantaishir serpentinite left). We favor the first possibility, as detailed below. mélange and the Dzabkhan–Baydrag microcontinent. This age and a similar metamorphic age (510 ± 4 Ma; Table 4) of a felsic granulite 7.5. Phase 5: continuing orogeny with local uplift in the Dzabkhan–Baydrag microcontinent (Kozakov et al., 2002)are interpreted here to represent the time of thrusting of the Dariv– The Gobi Altai microcontinent underwent crustal reworking in the Khantaishir ophiolite and the Lake arc over the Dzabkhan–Baydrag period ca. 518–498 Ma (Fig. 12F, right; Table 4; Hanžl and Aichler, microcontinent. Almost contemporaneously, low-Al TTG-like granitoid 2007; Hrdličková et al., 2008; Demoux et al., 2009b; Hrdličková et al., magmatism (ca. 519–494 Ma; Fig. 12D) and alkaline magmatism 2010; Kröner et al., 2010), suggested mainly by granitoids, some of

(ca. 511 Ma) developed in the Lake zone (Rudnev et al., 2009, 2012; which have near-zero and negative εNd(t) values (−1.5 to −0.2; Kovach et al., 2011; Yarmolyuk et al., 2011), while in the Dzabkhan– Hrdličková et al., 2010). This granitic episode closely followed on the Baydrag microcontinent there was broadly coeval granitoid magmatism generation of the Gobi Altai ophiolite (ca. 523–518 Ma), and we inter- at ca. 515–482 Ma (Kozakov et al., 2002; Dijkstra et al., 2006; Kröner pret it to reﬂect obduction of an ophiolite slice onto the Gobi Altai et al., unpublished, for analytical data see Table S2). microcontinent (Fig. 12E). The very short time span between such According to Rudnev et al. (2009, 2012) and Kovach et al. (2011) the SSZ-type ophiolite formation and obduction has been documented else- generation of low-Al TTG-like granitoids (εNd(t) =7.9–6.5; Table 4) where, such as Semail in Oman (e.g., Hacker and Gnos, 1997). This requires the melting of metabasic rocks at pressures of ca. 3–7kbarin implies that the subduction zone was initiated near the Gobi Altai the mid-upper crust (Fig. 12D). The heat source likely came from slab microcontinent (i.e. Fig. 12E, right). The existence of a subduction– delamination and consequent uprising asthenosphesre. The presence accretion complex with eclogite lenses in the Gobi Altai area (Fig. 3) of a near-coeval alkaline granite (511 ± 2 Ma; Table 4; Yarmolyuk provides geological evidence for this, although the timing of eclogite- et al., 2011)andaperidotite–pyroxenite–granodiorite suite (511 ± facies metamorphism has not been constrained. 7 Ma; Kovach et al., 2011) would be an expected outcome of such a The presence of an early Ordovician angular unconformity in Gobi process. The granitoids in the Dzabkhan–Baydrag microcontinent and Altai (Kröner et al., 2010) suggests that the Gobi Altai microcontinent a leucogranite (487 ± 7 Ma; Table 4; this work) cutting the Dariv was uplifted before the Ordovician, as in the Dzabkhan–Baydrag ophiolite have negative εNd (t) values (−10.5 to −12; Table 4; Kozakov microcontinent. This pre-Ordovician uplift was almost synchronous, et al., 2002; this work) indicating a predominant ancient crustal source. and accompanied extensive crustal reworking in the South Mongolia The broadly granitoid magmatism is thus interpreted to reﬂect crustal Gobi Altai (ca. 518–498 Ma; Fig. 12F, right) as well as in the Dzabkhan– reworking and anatexis, due to overthrusting and consequential crustal Baydrag microcontinent of Northwest Mongolia (ca. 515–482 Ma; thickening (Fig. 12D). Fig. 12D, right). In a classic collisional orogen, where an intervening During this phase, slab delamination, overthrusting and crustal ocean is closed and two formerly separated continents are coupled, reworking formed a continuum, which we ascribe to post-collisional surface uplift with crustal reworking marks the extensional collapse of tectonics. This continuum is thought to have accompanied surface uplift. the orogenic system that terminates the orogeny (e.g., Dewey et al., Evidence in support of this interpretation is the presence of a conglom- 1993; Rey et al., 2001; Jadamec et al., 2007). erate overlying the Lake arc that contains a 492 ± 1 Ma granite clast In our case, however, with surface uplift and crustal reworking in (Table 4; Kröner et al., 2001). Northwest Mongolia and South Mongolia Gobi Altai a huge accretionary complex to the south of the Gurvan Sayhan–Zoolen forearc (Fig. 12F, 7.4. Phase 4 (ca. 523–511 Ma): initiation of new subduction zone(s) in left) in the South Gobi zone (Fig. 1A)-Northern orogen of Chinese South Mongolia Inner Mongolia at ca. 498–415 Ma (Jian et al., 2008)wasbuilton a north-dipping thrust belt. The accretionary complex includes Following Phase 2 (ca. 535–524 Ma), i.e. arc-microcontinent colli- the following main units: (1) a large near-trench, TTG-like pluton sion in Northwest Mongolia, a new, north-dipping subduction zone (ca. 498–461 Ma; Chen et al., 2000; Jian et al., 2008); (2) a juvenile began to form in the Trans-Altai (Fig. 12E, left), South Mongolia. This plutonic arc (ca. 484–469 Ma; Jian et al., 2008); (3) a wide subduction– is indicated by the generation of the Gurvan Sayhan–Zoolen forearc accretion complex (Tang, 1990; Badarch et al., 2002) that underwent (520 ± 5–511 ± 5 Ma; Table 4; this work) with an ophiolite basement migmatization and anatectic melting in the period ca. 440–428 Ma (Shi Author's personal copy

90 P. Jian et al. / Earth-Science Reviews 133 (2014) 62–93 et al., 2003; Jian et al., 2010b); and (4) a blueschist-facies high-P/T Acknowledgments (pressure/temperature) metamorphic complex (Xu et al., 2001). This suggests an overall migration of the subduction zone southwards An earlier version of this paper was significantly improved by (Fig. 12F, left). The Gurvan Sayhan–Zoolen forearc with an ophiolite comments of the Journal reviewers (anonymous) and Editor (Manfred (Nca. 494 Ma) was likely uplifted at some time by continued frontal Strecker). The Natural Science Foundation of China (Grant nos. accretion (Fig. 12F, left), a mechanism suggested by Shervais 41030314 and 40234045) and the Geological Survey of China (Grant (2001) to explain the exposure of the Coast Range Ophiolite (CRO) nos. 1212010711817 and 1212010050504) supported this research. of California. Note also that along strike of the Trans-Altai zone the early to mid- Paleozoic Chinese Altai in Northwest China (Sun et al., 2009) and the References ca. 503–481 Ma Zhaheba ophiolite (Fig. 1A; Jian et al., 2003; Xiao Amantov, V.A., Blagonravov, V.A., Borzakovskiy, Y.A., Durante, M.V., Zonenshain, L.P., et al., 2006), are both younger than the latest Neoproterozoic Lake arc Luvsandanzan, B., Matrosov, P.S., Suyetenko, O.D., Filippova, I.B., Hasin, R.A., 1970. (ca. 551–540 Ma; Fig. 12B). This is broadly consistent with the southerly Main features of Paleozoic stratigraphy of Mongolian People's Republic. In: Luvsandanzan, B., Marinov, N.A., Zaitsev, N.S. (Eds.), Stratigraphy and Tectonics of migration of the subduction zones, as illustrated in Fig. 12F (left). More- the Mongolian People's Republic. Nauka Press, Moscow, pp. 8–63 (in Russian). over, comparable to the South Gobi zone-Northern orogen of Chinese Anon, 1972. Penrose field conference on ophiolites. Geotimes 17, 24–25. 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Mesozoic contact metamorphism in the western United States. In: Ernst, W.G. (Ed.), Metamorphism and Crustal Evolution of the Returning to the fundamental geological problem we earlier identi- Western United States (Rubey volume). Prentice-Hall, Englewood Cliffs, New Jersey, fied and discussed, our new zircon and geochemical data, along with re- pp. 110–178. evaluation of published geological evidence (e.g., Kröner et al., 2010), Bird, F., 1979. Continental delamination and the Colordo plateau. J. Geophys. Res. 84, 7561–7571. suggest that the long-accepted view that the CAOB in West Mongolia Black, L.P., Jagodzinski, E.A., 2003. Importance of establishing sources of uncertainty for is separated by the Main Mongolian Lineament (MML) into a north- the derivation of reliable SHRIMP ages. Aust. J. Earth Sci. 50, 503–512. western Early Paleozoic orogen and a southern Late Paleozoic orogen Black, L.P., Kamo, S.L., Allen, C.M., Aleinikoff, J.N., Davis, D.W., Korsch, R.J., Foudoulis, C., 2003. 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