Manuscript Details

Manuscript number EARTH_2017_327_R1

Title Origin of post-collisional magmas and formation of porphyry Cu deposits insouthern Tibet

Article type Invited review article

Abstract

The recent discovery of large porphyry copper deposits (PCDs) associated with Miocene (22–12 Ma) granitoidmagmas in the eastern section of the Paleocene-Eocene Gangdese magmatic arc in the Himalaya-Tibetan orogenicbelt raises new questions about the origin of water-rich (≥4.5 wt.%), oxidized (ΔFMQ 1–3) magmas in continentalcollisional settings and their mineralization potential. We review the literature and compile available data on whole rockand isotope geochemistry for Cenozoic igneous rocks from Tibet, and add new zircon Ce4+/Ce3+ and Ti-in-zirconthermometry data to better understand variations in oxidation state and thermal evolution of these suites, which arekey controls on Cu mineralization. Six distinct Cenozoic igneous suites are defined: Paleocene-Eocene syn-collisionalGangdese magmatic arc rocks (ΔFMQ = -1.2 to +0.8) (suite I), and five broadly contemporaneous Miocene suites. Adistinct change in magmatism along the length of the belt occurs at around 88°E in the Miocene suites: to the east,porphyry copper mineralization is associated with a moderately oxidized, high-Sr/Y granitoid suite (suite II, ΔFMQ =+0.8 to +2.9) with minor occurrences of transitional (hybrid) monzonitic (suite III) and trachytic rocks (suite IV; both withzircon Ce4+/Ce3+ > 50-100, EuN/EuN* = ~0.5, and ΔFMQ = ~+1 to +2). To the west of 88°E, trachytic volcanic rocks(suite V) are more voluminous but more reduced (zircon Ce4+/Ce3+ < 50, ΔFMQ <+1), and are associated withsparse, poorly mineralized high-Sr/Y granitoids (suite VI) which are moderately oxidized (zircon Ce4+/Ce3+ = 20–100,ΔFMQ = ~+1 to +3). The Miocene high-Sr/Y granitoids have many compositional and isotopic similarities to thePaleocene-Eocene Gangdese arc rocks, and are interpreted to have been derived by melting of the hydrated arc root,with minor mantle input. In contrast, the highly evolved isotopic signatures of the Miocene trachytic rocks, combinedwith deep seismic profiles and a xenolith-derived geotherm, suggest their derivation from the underthrust IndianProterozoic subcontinental lithospheric mantle (SCLM) or old fore-arc Tibetan SCLM during phlogopite breakdown attemperatures of ~1100°C. Based on published geophysical data and tectonic reconstructions, we develop a model thatexplains the origin of the various Miocene magmatic suites, their spatial differences, and the origin of related PCDs.Following the early stages of continental collision (Eocene–Oligocene), shallow underthrusting of the Indian continentallithosphere and subcretion of Tethyan sediments (including oxidized carbonates and possibly evaporites) undereclogite facies conditions promoted the release of aqueous fluids, which hydrated and oxidized the base of theoverlying Tibetan plate. This metasomatism rendered the Tibetan lower crust fusible and fertile for metalremobilization. During the mid-Miocene, the Indian slab steepened in the eastern sector (east of ~88°E). In thiseastern belt, deeply derived trachytic magmas were trapped in melt zones at the base of the Tibetan crust, andvariably mixed with the crustally-derived, high Sr/Y granitoid magmas. They may also have released water thatcontributed to fluid-fluxed melting of the lower crust, producing voluminous high-Sr/Y granitoid magmas, which wereassociated with significant PCD mineralization. Hybridization between the trachytic magmas and lower crustal partialmelts is indicated by intermediate isotopic compositions, enriched Cr and Ni contents, and high Mg# in someintermediate-to-felsic (56–70 wt. % SiO2) high-Sr/Y granitoids. Trapping of the trachytic melts in deep crustal meltzones explains the relatively small volumes of trachytic magmas erupted at surface in the east. In contrast, to the westof ~88°E, subduction of the Indian plate has remained flat to the present day, preventing incursion of hotasthenosphere. Consequently, cooler conditions in the deep Tibetan lithosphere resulted in limited crustal melting andthe production of only small volumes of high-Sr/Y granitic magmas. Trachytic melts ascending from the underthrustIndian or Tibetan plate were able to pass through the cooler lower crust and erupted in greater volume at surface,whereas only small volumes of high-Sr/Y granitoid magma were generated and are not associated with significantPCD mineralization.

Keywords high-Sr/Y granitoids, trachytic magmatism, Indian subduction geometry, water-fluxed melting, Gangdese, porphyry copper deposits

Corresponding Author Rui Wang

Corresponding Author'sInstitution

CSIRO

Order of Authors Rui Wang, Roberto Weinberg, William Collins, Jeremy Richards, Di-Cheng Zhu

Suggested reviewers Mark Allen, Massimo Chiaradia, Yongjun Lu, Yildirim Dilek

1

1 Origin of post-collisional magmas and formation of porphyry Cu deposits in southern

2 Tibet

3 Rui Wang1,2*, Roberto F. Weinberg3, William J. Collins4, Jeremy P. Richards5, and Di-cheng

4 Zhu1

5

6 1State Key Laboratory of Geological Processes and Mineral Resources, and School of

7 Scientific Research, China University of Geosciences, Beijing 100083, China

8 2CSIRO Mineral Resources, Perth, WA 6151, Australia

9 3School of Earth, Atmosphere and Environment, Monash University, Building 28, Clayton,

10 Victoria 3800, Australia

11 4The Institute for Geoscience Research (TIGeR), Department of Applied Geology, Curtin

12 University, GPO Box U1987 Perth WA 6845, Australia

13 5Harquail School of Earth Sciences, Laurentian University, Sudbury, Ontario, P3E 2C6,

14 Canada

15

16 The recent discovery of large porphyry copper deposits (PCDs) associated with Miocene

17 (22–12 Ma) granitoid magmas in the eastern section of the Paleocene-Eocene Gangdese

18 magmatic arc in the Himalaya-Tibetan orogenic belt raises new questions about the origin of

19 water-rich (≥4.5 wt.%), oxidized (ΔFMQ 1–3) magmas in continental collisional settings and

20 their mineralization potential. We review the literature and compile available data on whole

21 rock and isotope geochemistry for Cenozoic igneous rocks from Tibet, and add new zircon

22 Ce4+/Ce3+ and Ti-in-zircon thermometry data to better understand variations in oxidation state

23 and thermal evolution of these suites, which are key controls on Cu mineralization. Six

24 distinct Cenozoic igneous suites are defined: Paleocene-Eocene syn-collisional Gangdese

25 magmatic arc rocks (ΔFMQ = -1.2 to +0.8) (suite I), and five broadly contemporaneous

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26 Miocene suites. A distinct change in magmatism along the length of the belt occurs at around

27 88°E in the Miocene suites: to the east, porphyry copper mineralization is associated with a

28 moderately oxidized, high-Sr/Y granitoid suite (suite II, ΔFMQ = +0.8 to +2.9) with minor

29 occurrences of transitional (hybrid) monzonitic (suite III) and trachytic rocks (suite IV; both

30 with zircon Ce4+/Ce3+ > 50-100, EuN/EuN* = ~0.5, and ΔFMQ = ~+1 to +2). To the west of

31 88°E, trachytic volcanic rocks (suite V) are more voluminous but more reduced (zircon

32 Ce4+/Ce3+ < 50, ΔFMQ <+1), and are associated with sparse, poorly mineralized high-Sr/Y

33 granitoids (suite VI) which are moderately oxidized (zircon Ce4+/Ce3+ = 20–100, ΔFMQ =

34 ~+1 to +3).

35 The Miocene high-Sr/Y granitoids have many compositional and isotopic similarities to

36 the Paleocene-Eocene Gangdese arc rocks, and are interpreted to have been derived by

37 melting of the hydrated arc root, with minor mantle input. In contrast, the highly evolved

38 isotopic signatures of the Miocene trachytic rocks, combined with deep seismic profiles and a

39 xenolith-derived geotherm, suggest their derivation from the underthrust Indian Proterozoic

40 subcontinental lithospheric mantle (SCLM) or old fore-arc Tibetan SCLM during phlogopite

41 breakdown at temperatures of ~1100°C.

42 Based on published geophysical data and tectonic reconstructions, we develop a model

43 that explains the origin of the various Miocene magmatic suites, their spatial differences, and

44 the origin of related PCDs. Following the early stages of continental collision (Eocene–

45 Oligocene), shallow underthrusting of the Indian continental lithosphere and subcretion of

46 Tethyan sediments (including oxidized carbonates and possibly evaporites) under eclogite

47 facies conditions promoted the release of aqueous fluids, which hydrated and oxidized the

48 base of the overlying Tibetan plate. This metasomatism rendered the Tibetan lower crust

49 fusible and fertile for metal remobilization.

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50 During the mid-Miocene, the Indian slab steepened in the eastern sector (east of ~88°E).

51 In this eastern belt, deeply derived trachytic magmas were trapped in melt zones at the base

52 of the Tibetan crust, and variably mixed with the crustally-derived, high Sr/Y granitoid

53 magmas. They may also have released water that contributed to fluid-fluxed melting of the

54 lower crust, producing voluminous high-Sr/Y granitoid magmas, which were associated with

55 significant PCD mineralization. Hybridization between the trachytic magmas and lower

56 crustal partial melts is indicated by intermediate isotopic compositions, enriched Cr and Ni

57 contents, and high Mg# in some intermediate-to-felsic (56–70 wt. % SiO2) high-Sr/Y

58 granitoids. Trapping of the trachytic melts in deep crustal melt zones explains the relatively

59 small volumes of trachytic magmas erupted at surface in the east.

60 In contrast, to the west of ~88°E, subduction of the Indian plate has remained flat to the

61 present day, preventing incursion of hot asthenosphere. Consequently, cooler conditions in

62 the deep Tibetan lithosphere resulted in limited crustal melting and the production of only

63 small volumes of high-Sr/Y granitic magmas. Trachytic melts ascending from the underthrust

64 Indian or Tibetan plate were able to pass through the cooler lower crust and erupted in greater

65 volume at surface, whereas only small volumes of high-Sr/Y granitoid magma were

66 generated and are not associated with significant PCD mineralization.

67

68 Key words: high-Sr/Y granitoids, trachytic magmatism, Indian subduction geometry, water-

69 fluxed melting, Gangdese, porphyry copper deposits

70

71 Table of contents

72 1. Introduction

73 2. Geological setting

74 2.1 Tectonics

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75 2.2 Magmatism

76 2.3 Metallogenesis

77 3. Geodynamic evolution of the Himalayan orogen.

78 3.1 Onset of Indian-Asian collision

79 3.2 Indian plate shallow subduction

80 3.3 The nature of subducted Indian lithosphere

81 3.4 Present-day configuration of the collision zone

82 4. Six magmatic suite in the Gangdese belt

83 4.1 Paleocene-Eocene Gangdese arc igneous rocks

84 4.2 Eastern Miocene high-Sr/Y granitoids

85 4.3 Western Miocene high-Sr/Y granitoids

86 4.4 Western Miocene trachyte suite

87 4.5 Eastern Miocene trachyte suite

88 4.6 Eastern transitional monzonites

89 5. Trace elements in zircons

90 5.1 Trace element characteristics

91 5.2 Magma oxidation state

92 5.3 Magma temperature

93 5.4 Petrogenesis implication

94 6. Discussion

95 6.1 Paleogene magmatism

96 6.2 Post-Eocene Indian plate shallow subduction

97 6.3 Origin of Miocene trachytes

98 6.4 Tibetan or Indian lithospheric mantle melting as a source for trachytes?

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99 6.5 Fluid-fluxed melting and oxidation of Tibetan lower crust in the Miocene: the

100 origin of high-Sr/Y granitoids

101 6.6 Mixing model for Miocene high-Sr/Y magmas

102 6.7 Thermal structure of the Miocene Gangdese belt

103 6.8 Geodynamic model

104 6.9 Metallogenic implications

105

106 1. Introduction

107 Porphyry copper deposits (PCDs) are generally associated with oxidized and H2O-rich

108 magmas, typical features of magmatic arcs (Burnham, 1979; Candela, 1992; Richards, 2003;

109 Sillitoe, 2010). In island arcs and continental arcs, where porphyry deposits form, it is

110 generally thought that oxidized, sulphur-rich fluids released from subducting slabs migrate

111 into the asthenospheric mantle wedge, where they cause partial melting and mobilization of

112 metals (Richards, 2003; Audétat and Simon, 2012), and ultimately transfer these metals into

113 the crust. Recently, large porphyry copper deposits (PCDs) have been found in association

114 with post-collisional (Miocene; 22–12 Ma), high-Sr/Y granitoid plutons emplaced in the

115 eastern section of the Paleocene-Eocene Gangdese magmatic arc in the Himalayan-Tibetan

116 orogenic belt (Fig. 1; Hou et al., 2004, 2015; Yang et al., 2009, 2016; Lu et al., 2015; Wang

117 et al., 2014a, b, 2015a). These discoveries raise questions about the nature of magmatic and

118 metallogenic processes during continental collision.

119 In the Gangdese magmatic arc, the porphyry-related intrusions are coeval with a suite of

120 Miocene potassic volcanic rocks (24–8 Ma), which have been collectively termed

121 ultrapotassic volcanic rocks (UPVs) in the literature (Williams, 2000; Williams et al., 2001,

122 2004; Ding et al., 2003; Chung et al., 2005; Zhao et al., 2009; Zhou et al., 2010; Wang et al.,

123 2014c; Guo et al., 2013, 2015; Liu D et al., 2014, 2015, 2017; Xu et al., 2017). In this paper,

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124 we refer to this suite as trachytic. However, there is a puzzling difference in the spatial

125 distribution of the Miocene intrusive (high-Sr/Y granitoid) and trachytic volcanic suites along

126 the length of the Gangdese belt, with abundant PCD-hosting granitoids cropping out east of

127 ~88°E, but only a few poorly mineralized granitoids to the west (Hou et al., 2004; Zhao et al.,

128 2009; Li et al., 2011; Wang et al., 2014c; Yang et al., 2016). In contrast, trachytic volcanic

129 rocks are relatively common along the Gangdese belt west of 88°E, but rare to the east (Fig.

130 1).

131 A number of hypotheses have been proposed to explain the unusual origin of these

132 Miocene PCDs and their high-Sr/Y granitoid hosts (reviewed by Wang et al., 2015a, and

133 Yang et al., 2016). These hypotheses typically link the deposits to recycling of the

134 subduction-fertilized, deeper sections of the Paleocene-Eocene arc, and many ascribe the

135 magmas to either the remelting of the Tibetan lower arc crust (Chung et al., 2003, 2009; Hou

136 et al., 2004; Li et al., 2011) or metasomatized Tibetan lithospheric mantle (Lu et al., 2015),

137 while some argue for hybrid magmas from both sources (Wang et al., 2015a; Yang et al.,

138 2015, 2016). However, none of these models explain all the features (geochemistry, water

139 content, redox state) of these Miocene granitoids, and especially their relationship with

140 coeval trachytic volcanic rocks. Deep crustal and mantle xenoliths entrained by the Miocene

141 trachytes provide direct information regarding crust-mantle hybridization (Chan et al., 2009;

142 Liu CZ et al., 2011; Liu D et al., 2014; Wang et al., 2016), and suggest a link between the

143 origin of high-Sr/Y granitoids and the coeval Miocene potassic volcanic rocks (Wang et al.,

144 2017a).

145 The transition from subduction-related magmatism in the Paleocene to collisional

146 magmatism in the Miocene is accompanied by significant changes in geochemical and

147 isotopic (Sr-Nd-Hf-O) magmatic compositions (Wang et al., 2015a, b; Yang et al., 2016),

148 suggesting that underthrusting of the Indian plate was a major control on the nature of

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149 Miocene Gangdese magmatism. Following the Indian-Asian collision at ~55–50 Ma (Van der

150 Voo et al., 1999; de Sigoyer et al., 2000; Meng et al., 2012; Ding et al., 2016; Zhu et al., 2015,

151 2017), the subsequent magmatic quiescence in the late Eocene–Oligocene reflects shallow

152 angle underthrusting of the Indian continental margin (Guillot et al., 2008; Ji et al., 2009;

153 Ding et al., 2016). However, recent seismic studies reveal contrasting Indian plate subduction

154 geometry from west to east, with shallow underthrusting beneath the western Gangdese belt,

155 and steep underthrusting in the east (Zhao et al., 2010). By combining the location of a high-

156 velocity seismic anomaly corresponding to the subducted Indian plate in the deep mantle with

157 the palaeogeographic position of India, Replumaz et al. (2010) suggested that steep

158 subduction of India was initiated before 25 Ma. We suggest that the transition to steeper

159 subduction in the east should have occurred in the Miocene, leading to the opening of an

160 asthenospheric mantle wedge in the east but not in the west. Such differences impact the

161 tectono-thermal regime of the evolving collisional system, and the impact on magmatism

162 shown by the differences in the spatial distribution of intrusive and volcanic suites in the

163 Miocene.

164 Here, we review major aspects of the geodynamic setting of the Himalayan-Tibetan

165 orogen and the Miocene evolution of the Gangdese magmatic arc. We start with a brief

166 summary of the geological setting followed by an overview of the geodynamic aspects of the

167 orogen. We then use a compilation of 288 published geochemical and isotopic analyses to

168 distinguish between six Cenozoic igneous suites exposed in the Gangdese belt. We add new

169 titanium-in-zircon thermometry and zircon Ce4+/Ce3+data to constrain the temperatures and

170 redox states of these suites, and new plagioclase compositional data to constrain the

171 magmatic water contents. These features are combined with geodynamic reconstructions to

172 derive an integrated petrogenetic model for Miocene magmatism and mineralisation in the

173 Gangdese belt.

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174

175 2. Geological setting

176 2.1 Tectonics

177 The Himalayan-Tibetan orogen is composed of (from south to north) the Himalayas,

178 Lhasa terrane, Qiangtang terrane, and Songpan-Ganze complex, separated from each other by

179 the Indus-YarlungTsangpo, Bangong-Nujiang, and Jinsha River sutures, respectively (Yin

180 and Harrison, 2000; Zhu et al., 2013). The Lhasa terrane is divided into northern, central, and

181 southern Lhasa subterranes, bounded by the Shiquan River–Nam Tso Mélange zone and the

182 Luobadui–Milashan fault, respectively (Fig. 1; Zhu et al., 2011, 2013). The core of the Lhasa

183 terrane consists of Archean and Proterozoic crystalline basement (Zhu et al., 2011), which is

184 considered to have rifted from the Gondwana margin in the Late Triassic (Zhu et al., 2013; Li

185 et al., 2016). The Lhasa terrane is thought to have collided with the Qiangtang terrane to the

186 north in the Early Cretaceous (Kapp et al., 2005; Zhu et al., 2016). Northward subduction of

187 Neo-Tethyan oceanic lithosphere beneath its new northern margin, represented by the

188 accreted Lhasa terrane, began in the Late Triassic or Early Jurassic (Chu et al., 2006). Whole

189 rock Nd and zircon Hf isotopic compositions of the granitoid rocks in the Lhasa terrane

190 suggest an old and isotopically evolved central Lhasa subterrane with juvenile northern and

191 southern subterranes (Zhu et al., 2011; Hou et al., 2015).

192 The India–Asia collision started at ~55–50 Ma when the Greater India plate (Indian

193 continental margin; Ali and Aitchison, 2005) first collided with the Lhasa terrane (Meng et al.,

194 2012; van Hinsbergen et al., 2012; Zhu et al., 2015; Ding et al., 2016). The thicker and more

195 rigid Indian craton continues to subduct beneath the Lhasa terrane to the present day (Kind

196 and Yuan, 2010; Replumaz et al., 2010; Zhao et al., 2010). Seismic tomographic studies

197 indicate that the Indian continental lithosphere (100 to 200 km thick) extends northward

198 below the Tibetan plateau, where it is in direct contact with the base of the south Tibetan

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199 crust, and where the Tibetan plate sub-continental mantle lithosphere (SCLM) appears to

200 have been removed (Chung et al., 2009; Nábělek et al., 2009).

201

202 2.2 Magmatism

203 North-directed Neo-Tethyan subduction beneath southern Tibet produced voluminous

204 Jurassic–Cretaceous calc-alkaline magmatism in the Lhasa terrane (Harris et al., 1986; Wen,

205 2007; Mo et al., 2008; Lee et al., 2011; Wang et al., 2017b; Zhu et al., 2017). In contrast to

206 most Jurassic-Cretaceous igneous rocks that show typical continental arc features, a suite of

207 ~90–85 Ma charnockites with adakite-like features have been reported from the eastern

208 Gangdese belt (Wen et al., 2008; Zhang et al., 2010). These adakite-like rocks were

209 interpreted to have been derived from partial melting of the lower crust during “flat-slab”

210 subduction of the Neo-Tethyan ocean (Wen et al., 2008), or from the partial melting of a

211 subducted oceanic slab in a mid-ocean ridge subduction setting (Guan et al., 2010; Zhang et

212 al., 2010). A systematic geochronological study of Gangdese arc rocks reveals a magmatic

213 gap or quiescent period between ca. 80 and 70 Ma (Wen et al., 2008; Ji et al., 2009).

214 Afterwards, rollback of the Neo-Tethyan slab at ~69–53 Ma and possibly slab breakoff at

215 ~53–50 Ma triggered a magmatic flare-up (Kapp et al., 2007; Wen, 2007; Wang et al., 2015b;

216 Zhu et al., 2015), represented by extensive Paleocene–Eocene I-type intrusive rocks and

217 widespread Linzizong volcanic successions (Mo et al., 2008; Zhu et al., 2015, 2017; Fig. 1).

218 A magmatic gap or quiescent period from ~40–30 Ma was followed by emplacement of a

219 large number of small-volume calc-alkaline to alkaline intrusions and potassic (trachytic)

220 volcanic rocks in southern Tibet during the Oligo-Miocene (Ding et al., 2003; Hou et al.,

221 2009; Lu et al., 2015; Wang et al., 2015a, 2016, 2017a; Yang et al., 2016).

222

223

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224 2.3 Metallogenesis

225 Three episodes of porphyry-type mineralization are recognized in southern Tibet (Fig. 1):

226 Jurassic, Paleocene-Eocene, and Miocene.

227 The Middle Jurassic Xietongmen (Xiongcun) district in the middle of the Gangdese

228 magmatic belt is a large magmatic-hydrothermal centre (Tafti et al., 2009, 2014; Tang et al.,

229 2015; Wang et al., 2017b) that hosts the Xietongmen (No. I: 219.8 Mt @ 0.43% Cu, 0.51g/t

230 Au and 3.87g/t Ag) and Newtongmen (No. II: 388.9 Mt @ 0.32% Cu, 0.18g/t Au and 0.87g/t

231 Ag) deposits, and a few smaller Cu-Au prospects (e.g., Tangbai and Zemoduola). The

232 intrusions related to the ore-forming events are 176-171 Ma quartz diorite and granodiorite

233 porphyries (Tafti, 2011; Tang et al., 2015; Wang et al., 2017b).

234 Only two small deposits are known to have formed in the Paleocene–Eocene: the Sharang

235 porphyry Mo deposit (52.25 ± 0.31 Ma; 10 Mt @ 0.061% Mo; Zhao et al., 2014) and the Jiru

236 porphyry Cu deposit (49.2 ± 1.7 Ma; 41.9 Mt @ 0.43 Cu; Zheng et al., 2014).

237 The largest porphyry copper deposits formed in the eastern Gangdese belt (east of 88°E)

238 in the Miocene, and include the 16.4 ± 0.5 Ma Qulong porphyry Cu-Mo deposit (1,420 Mt @

239 0.5% Cu; Yang et al., 2009), the 14.7 ± 0.3 Ma Jiama porphyry Cu-Mo deposit (1,054 Mt @

240 0.44% Cu; Ying et al., 2014; Zheng et al., 2016), and the smaller Tinggong, Bangpu,

241 Tangbula, and Zhunuo porphyry Cu-Mo deposits (Hou et al., 2011; Wang et al., 2015a; Fig.

242 1). The ore-forming intrusions are granodiorite porphyries or granite porphyries with high

243 Sr/Y ratios (here termed high-Sr/Y granitoids), with ages between 21–13 Ma (Yang et al.,

244 2016). In the western part of the belt (west of ~88°E), Miocene plutons are sparse, and are

245 associated with only a few, small porphyry deposits (e.g., Zhunuo porphyry Cu deposit;

246 Zheng et al., 2007).

247

248

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249

250 3. Geodynamic evolution of the Himalayan-Tibetan orogen

251 3.1 Onset of Indian-Asian collision

252 The timing of initial Indian-Asian collision is important for understanding the evolution

253 of the Himalayan-Tibetan orogen. Based on the magmatic, metamorphic, biostratigraphic,

254 and paleomagnetic data, most workers agree that the onset of Indian continental subduction

255 below Asia occurred 55–50 Ma ago (de Sigoyer et al., 2000; Weinberg and Dunlap, 2000;

256 Meng et al., 2012; DeCelles et al., 2014; Zhu et al., 2015; Ding et al., 2016). However, others

257 have proposed ages ranging from ~70 to 35 Ma (e.g., Yin and Harrison, 2000; Aitchison et al.,

258 2007). Palaeomagnetic data indicate that the Lhasa terrane was at a latitude of 24°N when it

259 collided with Greater India at ~50 Ma (Meng et al., 2012). For the next 16 m.y., the southern

260 margin of Asia remained almost fixed while the Greater India plate subducted beneath it (soft

261 collision; Ali and Aitchison, 2005), causing shortening in the Himalaya and early uplift of the

262 Tibetan Plateau. By the end of the Eocene (~34 Ma), the thicker Indian craton made contact

263 with the Asian margin (hard collision: van Hinsbergen et al., 2012), increasing compressional

264 stress and initiating northward displacement of the southern margin of the Asian plate (Meng

265 et al., 2012).

266

267 3.2 Indian plate shallow subduction

268 The Indian plate is thought to have underthrust the Asian margin at a shallow angle in the

269 Eocene-Oligocene, as indicated by the following lines of evidence: (1) igneous rocks formed

270 between 50–41 Ma in the Gangdese arc on the southern Asian margin show more

271 heterogeneous and lower εHf(t) and εNd(t) values than early Cretaceous-Eocene arc igneous

272 rocks (Chung et al., 2005; Ji et al., 2009; Wang et al., 2015b), suggesting the involvement of

273 old crustal material not previously present in this part of the Tibetan plate, such as the Indian

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274 continental crust; (2) Indian plate upper crustal rocks in the Himalayan orogen (longitude

275 80°–95°E) show evidence for an early Eocene (48–45 Ma) medium-pressure amphibolite-

276 facies metamorphic event, suggesting underthrusting to depths of ~20–30 km beneath

277 southern Tibet (Guillot et al., 2008; Ding et al., 2016); and (3) the ages of Gangdese belt

278 magmatic zircons record a magmatic gap between ~41 and 30 Ma (Ji et al., 2009), which is

279 interpreted to represent cessation of Gangdese magmatism due to shallow subduction (Chung

280 et al., 2005; Rowley and Currie, 2006).

281

282 3.3 Initiation of steep subduction

283 Based on the studies of Replumaz et al. (2004), Negredo et al. (2007), and Replumaz et al.

284 (2010), the initiation of steep subduction can be estimated through combining the deepest part

285 of the high wavespeed anomaly in the deep mantle and the palaeoposition of the subduction

286 slab front. By combining the location of this anomaly with palaeogeographical positions of

287 India, Replumaz et al. (2010) suggest the India initiated steep subduction before 15 Ma.

288 The geological evidence of steep subduction comes from the discovery eclogite in the

289 central Himalaya around 88.5°E (the eastern Gangdese belt), and the only Lu-Hf date directly

290 from Arun garnet is 20.7±0.4 Ma (Corrie et al., 2010). The preservation of UHP

291 metamorphism requires steep subduction to permit the rapid return of UHP rocks to the

292 surface (Leech et al. 2005). Therefore, we suggest that the transition to steeper subduction in

293 the east should have occurred in the Miocene, which led to the opening of an asthenospheric

294 mantle wedge in the east but not in the west. This proposal is in line with the recent seismic

295 data. The dip angle of the northward Indian lithospheric subduction is increasing from the

296 west to east. The structure in the eastern Gangdese at the present is still steep subduction,

297 farther west in the Pamirs, it is apparent that subduction breakoff is occurring now (Lister et

298 al., 2008).

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299

300

301 3.4 The nature of subducted Indian lithosphere

302 In order to understand the influence of the underthrust material on the post-collisional

303 magmatic evolution of the Gangdese belt, it is necessary to consider the nature of the

304 continental material involved. In addition to crystalline and clastic sedimentary rocks (such as

305 limestone, mudstone, and chert; Phillips et al., 2013), evaporates and carbonates are reported

306 in the Neoproterozoic sequence, NW Himalaya, India (Singh et al., 2006). Related carbonates

307 are also reported in the Proterozoic stratigraphy of the Lesser Himalaya (Saha, 2013), and

308 large volumes of Late Jurassic sabkhas containing evaporitic sulphates and minor chlorides

309 were likely subducted. Such deposits occur widely on the southern margin of the former

310 Tethys Ocean, and are found from across the Arabian Peninsula to Iran (Leeder and Zeidan,

311 1977). In addition, Cretaceous-Tertiary (K-T) boundary evaporites are found in the Malatya

312 Basin on the Anatolide-Tauride plate of the Neo-Tethys Sea (Ayyıldız et al., 2015). Eocene

313 Tethyan evaporites were very likely to have existed, especially at the leading edge of the

314 subducting Indian plate where the sedimentary facies was shallow water and deposited during

315 the “Eocene maximum”, when corals extended to latitudes 51°N. Northern India was ~20-

316 30°S at this stage, probably close to the tropic of Capricorn. (Scheibner and Speijer, 2008).

317 Carbonate-bearing coesite eclogite also occurs in the Tso-Morari crystalline complex in

318 eastern Ladakh, India, suggesting that the northern margin of the Indian continent was

319 covered by carbonates and evaporites (Mukherjee et al., 2003; Johnston et al., 2011).

320 Evidence for subduction of these materials is important, because unlike crystalline and clastic

321 rocks, carbonates and sulphates are oxidants (Hattori, 2014), and their subduction could have

322 affected the oxidation state of the orogen.

323

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324 3.5 Present-day configuration of the collision zone

325 The Tibetan-Himalayan system is composed of three major parts: the Indian, Tibetan, and

326 Asian lithosphere, from south to north. A large number of seismic arrays have been operated

327 across large sections of the Tibetan plateau for over two decades to reveal the lithospheric

328 structure of the collision zone.

329

330 Indian lithosphere

331 There is abundant seismic evidence for subduction of the Indian plate below the southern

332 Tibetan Plateau. Seismic tomographic studies indicate that the Indian continental lithosphere

333 (100 to 200 km thick) dips northward below the Tibetan plateau, but that the extent of

334 underthrusting decreases from west (~31º N, ~85° E) to east (~ 30º N, ~91° E), with a NE-

335 directed convergence vector (Kumar et al., 2006; Li et al., 2008; Kind and Yuan, 2010; Zhao

336 et al., 2010; Shokoohi Razi et al., 2014; Liang et al., 2016).

337 The geometry and lateral continuity of the underthrust Indian plate lithosphere is debated,

338 limited by non-uniform seismic station coverage and the imprecision of existing tomographic

339 images (Liang et al., 2016). Receiver-function images (Kumar et al., 2006; Zhao et al., 2010)

340 and body and surface wave tomographic models (Nunn et al., 2014) suggest a west to east

341 increase in the dip-angle of the Indian plate lithosphere, a decrease of Indian plate

342 lithospheric thickness, and lack of Tibetan SCLM in the west (Fig. A1). The west–east

343 variability of P-normal velocities beneath the Himalayas and southern Tibet indicates that the

344 subducted Indian continental lithosphere is not homogeneous (Hearn et al., 2011). Fast

345 velocities (~8.4 km/s) were detected at a depth of ~90 km below the Tibetan plateau, which

346 are interpreted to correspond to localized formation of eclogite during underthrusting of

347 Indian lower crust in the Miocene (Huang et al., 2009; Shokoohi Razi et al., 2014). This

348 seismically fast material extends to the north of 32º N in western Tibet. In contrast, below

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349 eastern Tibet north–south-trending low-velocity anomalies are dominant (Liang et al., 2012).

350 Low-velocity anomalies in the upper mantle have been interpreted as evidence of

351 fragmentation of the Indian lithosphere (Liang et al., 2016). Three dimensional iso-surface

352 plots (Figure 8 in Liang et al., 2016) for the S-wave model reveal west-to-east variations of

353 Indian lithosphere (high-velocity anomalies), characterized by a shallowly dipping and

354 relatively intact lithosphere between longitude 85°E and 88°E, but a fragmented, steeply

355 dipping lithosphere between 88°E and 91°E. This fragmentation appears to connect with N–

356 S-trending rift faults and basins at surface in Tibet. Although there is some debate on the

357 detailed 3D structure of the subducting Indian lithosphere under Tibet at present, the main

358 feature revealed by geophysical data above is flat subduction occurring in the west and

359 steepening to the east.

360

361 Tibetan lithosphere

362 Lithosphere structure inferred from elevation, gravity and geoid anomalies, and

363 International Deep Profiling of Tibet and Himalaya (INDEPTH) surveys, reveal a Tibetan

364 lithosphere with thickness of ~180–200 km beneath central and northern Tibet that thins

365 southward where underlain by subducting Indian lithosphere (Kumar et al., 2006; Jiménez-

366 Munt et al., 2008; Zhao et al., 2011).

367 Seismic data indicate that the crust in southern Tibet is ~75 km thick, consisting of ~50

368 km of Tibetan crust underthrust by ~25 km of Indian crust (Owens and Zandt, 1997; Nábělek

369 et al., 2009). Magnetotelluric data from the Tibetan–Himalayan orogen from 77°E to 92° E

370 show an extensive low resistivity zone, interpreted to be a partially molten layer, along the

371 southern margin of the Tibetan plateau (Unsworth et al., 2005). Numerical models suggest

372 the possibility of channel flow of the partially molten layer under Tibet towards the

373 Himalayan front driven by lateral pressure gradients due to the topographic elevation

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374 differences (Clark and Royden, 2000; Beaumont et al., 2001). However, Miocene high-Sr/Y

375 granitoid magmas derived from the lower crust in southern Tibet (Hou et al., 2004) and

376 leucogranites in the High Himalayan Crystalline Series (Guo and Wilson, 2012) show very

377 different geochemical compositions, and do not support a connected, homogeneous lower

378 crustal melt sheet in the Miocene.

379 Based on the evidence summarized above, the Indian-Asian collision occurred at ~55–50

380 Ma, followed by shallow subduction of the Indian plate under the Asian margin. The

381 subduction angle remained flat in the western Gangdese belt, but steepened in the east in the

382 Miocene (Replumaz et al., 2010). Present-day configuration suggests that possible magmatic

383 sources for Miocene igneous rocks in southern Tibet are the Tibetan and Indian lithospheres,

384 and/or asthenosphere.

385

386 4. Six Cenozoic magmatic suites in the Gangdese belt

387 We have compiled published whole-rock geochemical data for 288 least-altered

388 Paleocene-Eocene (65–42 Ma) intrusive rocks and Miocene (24–8 Ma) volcanic and intrusive

389 rocks from the Gangdese belt in order to assess spatial differences in composition and

390 petrogenesis along the belt. Because some sample locations in the eastern Gangdese belt are

391 associated with porphyry-type alteration and mineralization, we excluded any samples

392 with >2 wt. % LOI or which were described as significantly altered. Four samples from mine

393 sites that showed extreme K/Na ratios were also excluded. Whole-rock geochemical and

394 isotopic data are listed in Tables 1 and A1 together with references. During the Oligocene

395 magmatic gap (Ji et al., 2009) only minor volumes of igneous rocks were formed, restricted

396 to the Mingze-Chenba area, where small quartz monzonite and quartz monzonite plutons

397 were emplaced at ~30 Ma (Zheng et al., 2012; Wang et al., 2014a). These have not been

398 included in the Miocene suites.

886887888889890891892893894895896897898899900901902903904905906907908909910911912913914915916917918919920921922923924925926927928929930931932933934935936937938939940941942943944

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399 Six Cenozoic igneous suites crop out along the Gangdese belt (Fig. 1): (1) a voluminous

400 Paleocene–Eocene Gangdese arc calc-alkaline suite (including intrusive rocks and Linzizong

401 volcanic rocks) that crops out along the length of the belt; and five Miocene suites that are

402 subdivided at a longitude of ~88°E: (2) a sparse western Miocene high-Sr/Y granitoid suite;

403 (3) a more voluminous eastern Miocene high-Sr/Y granitoid suite; (4) a voluminous western

404 Miocene trachytic suite (trachyandesitic to trachydacitic volcanic and subvolcanic rocks,

405 commonly called ultrapotassic volcanic rocks, or UPV); (5) a sparse eastern Miocene

406 trachydacitic (UPV) suite; and (6) a transitional (hybrid) Miocene monzonitic granitoid suite,

407 which outcrops only to the east of ~88°E. A summary of the main features of these six suites

408 and their zircon and plagioclase elemental compositions is given in Table 1 and discussed

409 below.

410

411 4.1 Paleocene–Eocene Gangdese arc igneous rocks

412 The first and oldest Cenozoic suite consists of voluminous Paleocene–Eocene I-type

413 intrusive and volcanic rocks (Linzizong volcanic successions; 67.7–42.5 Ma; references for

414 all ages are provided in Table A1) that cover nearly 50 % of the Gangdese arc (Fig. 1). Mafic

415 to intermediate intrusive rocks (gabbros and diorites) are mainly composed of plagioclase,

416 pyroxene, and amphibole, whereas granitoids are hornblende-bearing with minor biotite. The

417 Linzizong volcanic successions extend in a 1600 km-long, E–W-trending belt and consist of

418 calc-alkaline basaltic-andesitic lava flows, tuffs, and breccias, and dacitic to rhyolitic

419 ignimbrites. Silica values range from 48–76 wt. % SiO2. They are calc-alkaline to high-K

420 calc-alkaline (Fig. 2A, B), and are characterized by light rare earth element (LREE) and

421 large-ion lithophile element (LILE) enrichments, with depletions of Nb, Ta, P, and Ti (Fig.

422 3A, B), and relatively low Sr/Y (mostly < 40; Fig. 4A) and low La/Yb ratios (mostly < 30;

423 Fig. 4B). They have moderate to high εNd(t=15Ma) values (-4.4 to +8.5), low (87Sr/86Sr)t=15Ma

9459469479489499509519529539549559569579589599609619629639649659669679689699709719729739749759769779789799809819829839849859869879889899909919929939949959969979989991000100110021003

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424 ratios (0.7036–0.7068) (calculated at t = 15Ma to allow for direct comparison with Miocene

425 suites, Fig. 5A). Zircon εHft values range from -5.3 to +13, with mantle-like δ18O values

426 (+5.0 to +7.1 ‰; Wang et al., 2015b), and Nd-depleted mantle model (NdTDM) ages generally

427 <1 Ga (Table A1). Combined, the large isotopic and compositional ranges indicate that the

428 Paleocene–Eocene Gangdese arc rocks were a mix of mantle- and crustally-derived source

429 rocks.

430

431 4.2 Western Miocene high-Sr/Y granitoids

432 The rare western Miocene high-Sr/Y granitoids (18.4–16 Ma) are intermediate to felsic in

433 composition (granodiorite and quartz monzonite; Fig. 2A), with a restricted silica range (SiO2

434 63–70 wt. %). They are subalkalic and plot in the calc-alkaline to high-K calc-alkaline field

435 (Fig. 2B). They have higher Sr/Y ratios (44–99) and more significant LREE/HREE

436 fractionation (La/Yb = 26–51) than the Paleocene-Eocene suite, but their incompatible

437 element compositions are similar (Fig. 3A, B). The western high-Sr/Y granitoids have more

438 evolved Sr-Nd isotopic compositions than the Gangdese arc igneous rocks (Fig. 5A), with

439 εNd(t=15Ma) values from -9.3 to -4.1, and (87Sr/86Sr)t=15 Ma ratios from 0.7072 to 0.7100. Their

440 zircon εHft values range from -12.5 to +4.3, and δ18O values range from +6.2 to +8.1‰.

441 Their Nd depleted mantle model (NdTDM2) ages cluster between 1.0 and 1.4 Ga (Table A1).

442 These magmas were either derived from a crustal source slightly more evolved than the

443 Gangdese arc, or from a relatively homogeneous mix of Gangdese arc crust and a more

444 evolved source. Compositionally, they form a cluster between the more felsic components of

445 Gangdese arc magmas and the western Miocene trachyte suite (Figs. 2, 4).

446

447 4.3 Eastern Miocene high-Sr/Y granitoids

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448 The more abundant eastern Miocene high-Sr/Y granitoid suite (21.3–13.4 Ma) crops out

449 east of 88°E as shallowly-emplaced, small-volume plugs, dikes, or sills of porphyritic rocks,

450 locally associated with major porphyry copper deposits. The dominant lithologies are quartz

451 monzonite, granodiorite, and granite with a silica range mainly between 63–73 wt.% SiO2

452 (Fig. 2A). They are subalkaline and plot in the calc-alkaline to high-K calc-alkaline fields

453 (Fig. 2B). They have arc-like geochemical features with LILE enrichments and depletions in

454 Nb, Ta, and Ti (Fig. 3A, B), similar to the Paleocene–Eocene suite but with greater depletions

455 of HREE (Fig. 3B). These rocks are commonly referred to as high-Sr/Y granitoids (Chung et

456 al., 2003; Hou et al., 2004; Guo et al., 2007) because the great majority have Sr/Y ratios >50

457 (Fig. 4A) and high La/Yb ratios (26–80) (Fig. 4B). They have εNd(t=15Ma) values from -8.1 to

458 +5.7, and (87Sr/86Sr)t=15 Ma ratios from 0.7046 to 0.7082 (Fig 5A), overlapping with their

459 western counterparts but extending to more primitive compositions typical of the Paleocene–

460 Eocene suite. Their Nd depleted mantle model (NdTDM2) ages range from 0.3–1.3 Ga (Table

461 A1). They have more mantle-like zircon εHft (+1.4 to +8.7) and zircon δ18O values (+5.5 to

462 +6.5 ‰) compared to their western equivalents, and the range to more primitive Sr and Nd

463 isotopic compositions suggest they have a large component of mantle-derived Gangdese arc

464 crust as a source component (Fig. 5A,B).

465

466 4.4 Western Miocene trachyte suite

467 Miocene alkaline (trachytic) volcanic rocks (Fig. 2A) crop out mostly in the western

468 Gangdese belt where they are locally related to N–S trending grabens or normal faults (Cogan

469 et al., 1998; Williams et al., 2001, 2014; Lu et al., 2017); they are relatively rare in the east

470 (see below). The volcanic rocks are porphyritic with phenocrysts mainly of olivine,

471 clinopyroxene, phlogopite, and sanidine in a glassy or fine-grained groundmass. They plot as

472 trachyandesite (latite) and trachydacite on a total alkali–silica (TAS) diagram (Fig. 2A) and

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473 most analysed samples cluster between 53–71 wt.% SiO2 (Fig. 2A). Their high K2O contents

474 place them in the shoshonite (SH) field (Fig. 2B), or the ultrapotassic field according to Foley

475 et al. (1987). Based on the latter definition, more than 50% of the southern Tibet samples are

476 ultrapotassic (Fig. 2C, D), and this suite of rocks has widely been referred to as ultrapotassic

477 volcanic rocks, or UPV (Williams, 2001, 2014; Zhao et al., 2009; Liu et al., 2014; Wang et

478 al., 2015a); however, we prefer the IUGS-consistent term “trachytic” (Le Bas and Streckeisen,

479 1991). The mafic end of this suite has 10–12 wt. % MgO at ~52–55 wt. % SiO2 (Fig. 4E), and

480 the intermediate compositions are similar to high magnesium andesites (Wood and Turner,

481 2009), indicating that they are primitive, mantle-derived magmas. The western trachytes have

482 higher LREE contents than most of the Miocene granitoids (Fig. 3C, D), but similarly strong

483 negative Nb-Ta anomalies. The more mafic rocks are particularly rich in Cr (up to 649 ppm)

484 and Ni (up to 467 ppm), both of which show a strong positive correlation with Th (Fig. 4F).

485 The more felsic (trachydacitic) rocks have compatible element and HFSE contents similar to

486 the high-Sr/Y granitoids. They show variable Sr/Y (11–113) and La/Yb ratios (31–211),

487 which overlap with the lower range of the Paleocene-Eocene suite, but extend to much higher

488 ratios characterized by the high-Sr/Y granitoids. The trachytic suite is also characterized by

489 high (87Sr/86Sr)t=15 Ma ratios (0.7069–0.7393), strongly negative εNd(t=15Ma) values (-18.5 to -

490 7.1) (Fig. 5A), high NdTDM2 values (1.2–2.5 Ga, Table A1), mostly negative zircon εHft

491 values (-14.7 to +1.0), and crust-like zircon δ18O values (+6.9 to +8.3 ‰), all indicative of an

492 old lithospheric source. This crust-like isotopic range contrasts with their mantle-like

493 (187Os/188Os)t values (0.156–0.188) and high Cr-Ni contents (Wang et al., 2015a), and is

494 interpreted to reflect an ancient subcontinental lithospheric mantle (SCLM) source for the

495 high MgO (10–12 wt.%) end-member.

496

497 4.5 Eastern Miocene trachyte suite

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498 Upper Miocene (12–11 Ma) trachytic volcanic rocks are only known to occur in two

499 locations in the eastern Gangdese region, at Yangying and Suojin (Fig. 1). They are

500 essentially similar to the western trachytes, but are more silicic, with lower CaO and MgO

501 contents, and lower Mg# (Fig. 4). They are classified on the TAS diagram as trachydacites

502 (Fig. 2A) with K2O contents above 3 wt. % (Fig. 2B), and plot in the shoshonitic field in

503 Figure 2B. Phenocrysts include clinopyroxene, amphibole, plagioclase, and phlogopite, set in

504 a fine-grained groundmass (Wang et al., 2015a). Like the western trachytes, these

505 trachydacites are characterized by high (87Sr/86Sr)t ratios (0.7119–7121), strongly negative

506 εNdt values (-9.8 to -9.2), variable zircon εHft values (-5.7 to +10.4), intermediate zircon

507 δ18O values (+5.0 to +6.7 ‰), and mantle-like (187Os/188Os)t values (0.153–0.210) (Wang et

508 al., 2015a). They overlap in whole rock and isotope chemistry with the western trachytes, but

509 their uniformly low MgO, Ni and Cr contents suggests they are more fractionated equivalents

510 of those rock types.

511

512 4.6 Eastern transitional monzonites

513 The sixth igneous suite is transitional between the trachytes and high-Sr/Y granitoid rocks,

514 and has ages between 16–14 Ma. They occur in the ore fields such as Qulong, Jiama, and Jiru

515 mine in the eastern Gangdese belt, but they are not the ore-forming magmas in these deposits.

516 They are post-mineralization, and cross cut ore-forming granitoids (high-Sr/Y granitoids).

517 The Qulong transitional rocks have been investigated by Yang et al. (2015), who called them

518 high-Mg diorites; however, they plot as monzonites in Figure 2A, and in the high-K calc-

519 alkaline and shoshonitic fields in Figures 2B and 2D, respectively. They are silica

520 oversaturated, with a primary mineral assemblage of plagioclase, quartz, amphibole, and

521 minor biotite. Although silica contents range between of 59–63 wt. %, they have high K2O

522 (3.0–3.8 wt. %), relatively high MgO (4.1–5.4 wt. %) and Mg# (56 and 66), and high

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523 CaO/Al2O3 ratios (0.30–0.36) (Fig. 4D). Their Mg# values are much higher than the high-

524 Sr/Y granitoids (<50; Fig. 5A), but their LILE and HFSE contents, including high Sr/Y ratios

525 (Fig. 4A), are similar to many of the eastern granitoids. Accordingly, we refer to these as

526 “transitional” monzonites. Their transitional character is also evident in their isotopic

527 characteristics. They have low to moderate (87Sr/86Sr)t=15 Ma ratios (0.7057–0.7072),

528 moderately negative εNd(t=15Ma) values (-6.5 to -3.4), low (187Os/188Os)t ratios (0.176–0.178),

529 highly variable zircon εHft values (-1.1 to +7.1), and low mantle-like zircon δ18O values

530 (+5.3 to +6.1 ‰). Thus, the transitional monzonites have affinities to both the trachytes and

531 eastern Miocene high-Sr/Y granitoids (Figs. 2–5), suggesting melting in the deep crust, but

532 with an additional component of trachytic magma.

533

534 5. Trace elements in zircons

535 We have compiled trace element analyses of zircon from the Gangdese Cenozoic

536 magmatic suites from the literature, and added 44 new analyses from 7 new samples to

537 constrain magmatic conditions. Analytical methods are described in the Appendix and the

538 data are listed in Table A2.

539

540 5.1 Trace element characteristics

541 Trace elements in zircon reflect the characteristics of the magmas from which they

542 crystallized (Ballard et al., 2002; Rubatto, 2002; Trail et al., 2012; Kirkland et al., 2015).

543 Here, we are particularly interested in indicators of magmatic oxidation state, temperature,

544 and evolution, and so focus on zircon Ce4+/Ce3+ ratios, Ti contents, and trace element patterns.

545 Ce4+/Ce3+ ratios have been used to distinguish between relatively oxidized ore-bearing

546 porphyries (Ce4+/Ce3+ mostly > 50) from more reduced, barren intrusive suites (Ce4+/Ce3+ <

547 50; Ballard et al., 2002). EuN/EuN* ratios (where EuN/EuN* = EuN/(SmN*GdN)) in zircons

12401241124212431244124512461247124812491250125112521253125412551256125712581259126012611262126312641265126612671268126912701271127212731274127512761277127812791280128112821283128412851286128712881289129012911292129312941295129612971298

23

548 also correlate with oxidation state because Eu2+ is excluded from zircon relative to Eu3+.

549 Interpretation of magmatic redox state from the EuN/EuN* ratios in zircons is complicated by

550 the effects of plagioclase crystallization, which preferentially partitions Eu2+ relative to Eu3+.

551 However, in water-rich magmas, plagioclase crystallization is delayed until late in the

552 crystallization history (Naney, 1983), so should have minimal effect on zircon EuN/EuN*

553 ratios in hydrous, intermediate composition rocks (Dilles et al., 2015).

554 Zircon Th/U ratios can also be used to assess the degree of crystal fractionation (Kirkland

555 et al., 2015), and titanium (Ti) concentration can be used to estimate magmatic crystallization

556 temperatures (Ti-in-zircon; Watson and Harrison, 2005). Figure 6 reports the mean values

557 and standard deviation of zircons from each rock sample, and Fig. A2 illustrates single spot

558 analytical results for each zircon grain investigated here.

559

560 5.2 Magmatic oxidation state

561 Paleocene-Eocene Gangdese arc rocks and the western trachyte suite are characterized by

562 low zircon Ce4+/Ce3+ (mostly <50) over a wide spectrum of EuN/EuN* ratios (mostly <0.5)

563 (Figs. 6B-C, A2A, Table 1). The western high-Sr/Y granitoids have higher zircon Ce4+/Ce3+

564 (up to 159) and EuN/EuN* ratios (up to 0.78). The eastern Miocene high-Sr/Y granitoids show

565 similar zircon Ce4+/Ce3+ and EuN/EuN* ratios to the sparse western suite, but extend to

566 slightly higher values in EuN/EuN* (up to 0.87).

567 The Ce4+/Ce3+ ratios can be used to indicate relative magmatic oxidation state, and the

568 results reported here are consistent with a previous study of magnetite-ilmenite mineral pairs

569 which showed that Paleocene-Eocene Gangdese arc rocks have low to moderate ΔFMQ

570 values (-1.2 to +0.8; where ∆FMQ is measured in log fO2 units relative to the fayalite–

571 magnetite–quartz oxygen buffer), whereas eastern Miocene high-Sr/Y granitoids are more

572 oxidized (ΔFMQ+0.8 to +2.9; Wang et al., 2014b).

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573 Zircons from the eastern “transitional” monzonite suite show exceptionally high zircon

574 Ce4+/Ce3+ ratios (up to 487; Fig. 6C, Table A2), but their EuN/EuN* ratios are not

575 correspondingly high (~0.5). The sparse eastern trachytes also have elevated Ce4+/Ce3+ and

576 EuN/EuN* ratios, but hot as high as the western trachytes. The decoupling of zircon Ce4+/Ce3+

577 and EuN/EuN* ratios in the eastern transitional monzonites and trachytes suggest their

578 oxidation states are not extremely high, likely close to ΔFMQ +1 to +2 (Wang et al., 2014b;

579 Lu et al., 2016).

580 In summary, Ce4+/Ce3+ and EuN /EuN* ratios in zircons show that a spatial and temporal

581 distribution of oxidation state exists along the length of the Cenozoic Gangdese belt.

582 Paleocene-Eocene and western Miocene suites are relatively reduced with low Ce4+/Ce3+ and

583 scattered EuN/EuN* ratios, but all three eastern Miocene suites are relatively oxidized with

584 Ce4+/Ce3+ ratios >50. These observations suggest that oxidation of the deep Tibetan

585 lithosphere occurred after Paleocene–Eocene magmatism ceased and continental collision

586 began.

587

588 5.3 Magmatic temperature

589 The titanium-in-zircon geothermometer depends on the activity of Si and Ti in the host

590 magma. Overestimation of aSiO2 or aTiO2 yields overestimates and underestimates of

591 temperature, respectively, but values are generally considered to be correct within ±50 °C if a

592 Ti-bearing phase is present (McDowell et al., 2014). Rutile occurs in the trachytic rocks,

593 titanite in the Gangdese arc rocks, and ilmenite and/or titanite in the high-Sr/Y granitoids, so

594 that magma temperature estimates are thought to be reasonably accurate. Only some of the

595 more mafic Gangdese arc rocks and most of the trachyte suite are likely to have

596 underestimated temperatures because of high Zr solubility in mafic and alkaline magmas

597 (Watson & Harrison, 1983).

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598 Almost all calculations indicate temperatures <900°C. The Paleocene-Eocene Gangdese

599 arc zircons show a progressive decrease in temperature from ~870°C to 600°C as a function

600 of decreasing Th/U (Figs. 6D, A2B), reflecting progressive fractional crystallization to

601 subsolidus conditions (Kirkland et al., 2015). Zircons in the eastern high-Sr/Y granitoids have

602 temperatures typically <800°C and no obvious trend with Th/U values (Fig. A2B), suggesting

603 crystallisation from cooler magmas. Data from the sparse western high-Sr/Y granitoids

604 overlap this range but extend to higher temperatures (up to 818°C) and Th/U ratios (up to 2.6)

605 (Fig. A2B), suggesting crystallization from somewhat hotter melts than in the east. The

606 trachytic rocks generally record higher temperatures than in the granitoid suites (980°C to

607 700°C; Figs. 6D and A2B), as expected for mantle-derived magmas.

608

609 5.4 Petrogenetic implications

610 Chondrite-normalized zircon multi-element patterns are mostly steep for all magmatic

611 suites, but some of the western trachytic rocks have flatter HREE patterns (Fig. 6A),

612 particularly those trachytes with high Sr/Y ratios and large La/Yb variations. It is likely that

613 these flat HREE patterns are a result of partial melting or crystallization in equilibrium with

614 garnet (Rubatto, 2002). Such patterns are not observed in the high-Sr/Y granitoids, even

615 though their high Sr/Y ratios might be taken to indicate melting in the garnet stability field

616 (Macpherson et al., 2006). Instead, this lack of evidence for garnet fractionation in the zircon

617 REE patterns, along with relatively low whole-rock La/Yb ratios, suggest that early

618 amphibole fractionation and delayed plagioclase fractionation from hydrous melts was

619 responsible for the observed trace element characteristics (Richards and Kerrich, 2007;

620 Richards, 2011; Wang et al., 2014a).

621 The Ce4+/Ce3+ vs. T(ti-zr) (Ti-in-zircon) plot (Fig. 6E) shows two key features: (1) the

622 Miocene high-Sr/Y granitoids have higher Ce4+/Ce3+ ratios than most western trachytic rocks

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623 and Gangdese arc magmatic rocks; and (2) all Miocene granitoids show low T(ti-zr) below

624 750 °C. One eastern trachydacite sample records relatively high oxidation state

625 (exceptionally high zircon Ce ratios but intermediate EuN/EuN*) at high temperature,

626 implying that the oxidation state of those magmas is not controlled by crystal fractionation, as

627 could be implied from Fig. 6E. This disconnection with fractionation is further illustrated by

628 the Ce4+/Ce3+ vs Th/U plot (Fig. 6F), which shows that oxidation state is independent of Th/U

629 ratios for all the high-Sr/Y and transitional (monzonitic) granitoids. The data indicate that low

630 and high temperature magmas can be associated with high oxidation state, independent of

631 fractionation (Fig. 6F), and that the high-Sr/Y granitoids and eastern trachytic suites have

632 higher oxidation states than the Paleocene–Eocene and western trachytic suites.

633

634 6. Discussion

635 Miocene high-Sr/Y granitoids bear many similarities to the Paleocene-Eocene Gangdese

636 arc magmatic rocks, but crucial lithogeochemical and isotopic differences also suggest

637 affinities with the Miocene trachytic suites. We hypothesize that in the Miocene, alkaline

638 magmas hybridised with magmas generated by melting of the Gangdese arc root. Here, we

639 review possible petrogenetic models, then discuss tectonic settings that could explain this

640 magmatic evolution, as well as the implications for porphyry copper mineralization.

641

642 6.1 Paleogene magmatism

643 The onset of Paleocene magmas in the Gangdese belt has been ascribed to Neo-Tethyan

644 slab rollback (Chung et al., 2005; Ji et al., 2009; Lee et al., 2009; Wen et al., 2008; Zhu et al.,

645 2015). This is supported by southward migration of arc magmatism, an abrupt change of

646 India–Asia convergence between ~69–58 Ma (Lee and Lawver, 1995), and the development

647 of extension setting in the Qiangtang terrane characterized by east–west-trending sedimentary

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648 basins (Chung et al., 2005).

649 The igneous rocks during this time (~69–58 Ma) have relatively homogeneous and

650 juvenile isotopic compositions (εNdi of -0.6 to +4.0, εHfi of +3.8 to +7.1, and δ18O of +5.0 to

651 +6.5 ‰; Wang et al., 2015b), suggesting a significant juvenile crustal component mixed with

652 a minor (if any) mantle component (Wen et al., 2008; Wang et al., 2015b). Significantly, the

653 juvenile signatures highlight the absence of ancient Tibetan SCLM beneath the Gangdese arc

654 throughout this interval.

655 Changes in Sr-Nd-Hf-O isotopic compositions for the Linzizong volcanic rocks and coeval

656 intrusions in the Gangdese belt at ~53–50 Ma suggest mantle input and extensive crustal

657 melting, possibly associated with Neo-Tethyan slab breakoff (Wen et al., 2008; Wang et al.,

658 2015b; Zhu et al., 2015). The conclusion of slab breakoff at 53–50 Ma comes from many

659 lines of evidence, which include: I) asthenospheric influx was triggered by slab breakoff. The

660 Nd-Hf isotopic data of ~53–49 Ma igneous rocks show significant input from asthenospheric

661 mantle with εNdi values up to +9.8 and zircon εHfi values up to +15.1, which are comparable

662 with arc rocks from early periods (Wang et al., 2015b; Zhu et al., 2015); II) Voluminous

663 magmatism emplaced during ~53–49 Ma. It is suggested by widespread outcrops of

664 intrusions and Pa’na volcanic sequence at ~53–49 Ma (Fig. 2, Mo et al., 2003; Chung et al.,

665 2005; Lee et al., 2009); III) anomalously high magmatic temperatures (Tzr up to 800°C and

666 T(ti-zr) up to 980°C), reflecting a thermal anomaly at that time; IV) Bimodal volcanic rocks

667 have been reported in the Pa’na and Nianbo formations of Linzizong volcanic rocks (Mo et

668 al., 2003; Lee et al., 2009, 2011; Zhu et al., 2015); V) This high temperature event (or

669 thermal anomaly) led to extensive crustal melting, and generated heterogeneous magmatic

670 geochemistry during ~53–49 Ma. This is evidenced by heterogeneous whole-rock

671 geochemical compositions, heterogeneous zircon Hf isotopic compositions (εHf(t) = -5.3 to

672 15.1)), and scattered magmatic temperatures; and VI) crustal deformation, characterized with

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673 peak granulite-facies metamorphism from 66 to 52 Ma in the lower crust (Zhang et al., 2013).

674 These early Paleogene rocks are associated with few porphyry-type deposits, possibly

675 because the magmas were relatively anhydrous (Wang et al., 2014a) and less oxidized

676 (ΔFMQ -1.2 to +0.8) than magmas typically associated with PCDs (ΔFMQ +2) (Wang et al.,

677 2014b).

678

679 6.2 Post-Eocene flat Indian plate subduction

680 Crustal mass balance estimations suggest large-scale subduction of Indian continental

681 crust during the India-Asia collision (Ingalls et al., 2016; Capitanio et al., 2010). The

682 magmatic gap in the Gangdese belt between 40–30 Ma (section 3.2), and the gap in the

683 histogram of U-Pb ages for zircons from crustal xenoliths in “ultrapotassic” trachytic rocks in

684 the Himalayas (Liu et al., 2014), are consistent with flat subduction of the Indian plate

685 throughout the Oligocene. This suggests the Indian continent was subducted below southern

686 Tibet at a relatively shallow angle with little or no asthenospheric mantle wedge above the

687 underthrust Indian lithosphere (Guillot et al., 2008; Ding et al., 2016).

688 This magmatic gap was followed at ~35 Ma by a jump to higher DyN/YbN and U/Yb ratios

689 in xenocrystic zircons from Miocene trachytic rocks, suggesting they formed in thickened

690 crust, in equilibrium with eclogite and garnet-bearing granulitic mineral assemblages (Liu et

691 al., 2014). Moreover, the jump to negative zircon εHf values indicates a major change to an

692 isotopically evolved source for the trachyte suite at this stage. This is consistent with a recent,

693 more comprehensive study of magmatic zircons from Tibet (Liu et al., 2017), which also

694 shows an isotopically evolved source component added to the high-Sr/Y magmas between

695 32–28 Ma.

696 The underthrust Indian plate would have progressively dehydrated during continental

697 collision (Massonne, 2016). Underthrust crustal rocks, including some sediments, are likely

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29

698 sources of a hydrous fluid phase capable of metasomatising and/or melting the base of the

699 lower Tibetan crust. In particular, Eocene evaporites and carbonates were common in the

700 Neo-Tethys ocean, extending from Europe to SE Asia, and were prominent in the Great Kavir

701 Basin of Iran (Mukherjee et al., 2003; Johnston et al., 2011). Extensive carbonates also

702 existed on the leading edge of the Indian plate passive margin (Scheibner and Speijer, 2008).

703 They presumably also existed on the leading edge of the Indian plate passive margin, and

704 subduction of such oxidized materials could have affected the oxidation state of the Gangdese

705 arc root. Water released during flat subduction could have affected the rheology and changed

706 the mineralogy of the overriding lithospheric mantle, as revealed by high S-wave velocities

707 (Sommer and Gauert, 2011; Wagner et al., 2005). When flat subduction ends, a hot

708 asthenospheric mantle wedge opens between the two plates, which can cause partial melting

709 of the hydrated and oxidized upper plate lithosphere, as has been proposed for the central

710 Andes in the Miocene following a period of flat subduction (James and Sacks 1999; Kay et

711 al., 1999; Kay and Mopodozis, 2001). We suggest a similar development for the Gangdese

712 belt where initial flat subduction of the Indian plate was followed by steepening in the east

713 (Kumar et al., 2006; Kind and Yuan, 2010; Zhang et al., 2010), allowing for voluminous

714 high-Sr/Y granitoids to form in eastern Tibet. We will return to this idea when discussing the

715 Miocene geodynamic evolution of the region in section 6.8.

716

717 6.3 Origin of Miocene trachyte suites

718 Several single-source models have been proposed for the origin of the trachyte suite

719 (UPVs) in southern Tibet, including: (1) partial melting of the middle-lower crust of the

720 Indian plate (Hébert et al., 2014); (2) melting of enriched Asian (Tibet) lithospheric mantle

721 during delamination or convective thinning (Miller et al., 1999; Liu C et al., 2011), or due to

16531654165516561657165816591660166116621663166416651666166716681669167016711672167316741675167616771678167916801681168216831684168516861687168816891690169116921693169416951696169716981699170017011702170317041705170617071708170917101711

30

722 hydration by fluids from Indian plate subduction (Yang et al., 2016); or (3) derivation from

723 metasomatized asthenospheric mantle (Guo et al., 2013).

724 Several lines of evidence suggest these volcanic rocks were derived from an ancient

725 lithospheric mantle rather than crustal source: firstly, they carry not only lower-crustal but

726 also mantle xenoliths (Liu C et al., 2011; Liu D et al., 2014; Wang et al., 2016); secondly,

727 they have low SiO2 contents (down to 45 wt.%), high Mg# (up to 76), and high Ni and Cr

728 contents (467 and 649 ppm, respectively); and thirdly they have low, mantle-like Os isotopic

729 compositions (187Os/188Os)i = 0.154–0.210; Wang et al., 2015a). Indeed, the 10–12 wt. %

730 MgO contents at 55–60 wt. % SiO2 (Fig. 4E) require that these rocks were primary mantle-

731 derived magmas (Grove et al., 2012). Most significantly, Figure 7C–F shows that the most

732 evolved εNdi values are found in those rocks with highest values of MgO, Cr and Th, and

733 lowest values of SiO2. As silica increases and MgO decreases, the rocks become more

734 isotopically primitive (Fig. 7C, D). This observation is further supported by the low εNdi

735 ratios which, for the most MgO-rich endmembers, reach -18.5 (Fig. 5A) and have the highest

736 HFSE and LILE element contents (Fig. 7E,F). The Nd model ages ranging up to 2.5 Ga (Fig.

737 5D) indicate a probable Paleoproterozoic to latest Archean source.

738 A Tibetan SCLM origin for the trachytes was suggested by Yang et al. (2016). They

739 considered that devolatilization of the subducting Indian crust would have metasomatized the

740 overlying wedge of subcontinental lithospheric mantle, ultimately producing ultrapotassic

741 and/or alkaline mafic magmas. The major problem with this model is that no evidence exists

742 for Tibetan SCLM beneath the Gangdese arc during the Paleocene or Eocene. Accordingly,

743 the mantle wedge should have been asthenospheric and the resultant trachytic magmas should

744 be much less evolved. However, it is possible that ancient Tibetan SCLM was underthrust

745 from the forearc region beneath the Gangdese arc during Oligocene collision, and subsequent

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31

746 devolatilization during ongoing subduction would have resulted in progressive fluid-fluxed

747 melting of the metasomatized SCLM.

748 Guo et al. (2015) focused on the chemical and isotopic nature of the most primitive of

749 these post-collisional ultrapotassic (trachytic) magmas with MgO > 6 wt. %. These only crop

750 out west of 87°E. They proposed a two-stage model: During the first-stage (55–25 Ma),

751 fluids and melts released from the subducting Indian crust resulted in the formation of

752 pyroxenites through metasomatism of the overlying mantle wedge. During the second-stage

753 (25–8 Ma), partial melting of pyroxenites caused by slab roll-back and slab break-off

754 generated the trachytic magmas. They argued that the absence of these primitive endmembers

755 of Miocene ultrapotassic magmatism east of 87°E indicates different subduction geometries.

756 The critical problem of this model is that Guo et al (2015) assume that the decreases in Sri

757 and Pbi, and increasing Ndi of the Miocene rocks from the western to eastern Gangdese belt,

758 result from an eastward-decreasing crustal component in the mantle source region. However,

759 as we have shown above (Figs. 5 and 7), these changes are the result of an increasing crustal

760 component to the east, because the crustal (Gangdese arc) component in the high-Sr/Y

761 magmas is more juvenile than the mantle component in the trachytic magmas (see section

762 6.6).

763 Wang et al. (2016, 2017a) showed that the εNdi values for southern Tibet trachytes

764 increase with increasing SiO2. The positive slope is more likely to be part of a mixing array

765 between mafic trachytic magmas derived from an isotopically evolved source with felsic

766 magmas derived from the isotopically unevolved southern Tibetan (Gangdese) crust, as

767 suggested by xenolith evidence (Wang et al., 2017a).

768 Hybrid origins have also been proposed to explain the trachytic (UPV) suite. For example,

769 direct melting of subducted (Indian) crustal rocks followed by interaction with mantle

770 peridotite has been proposed to explain the origin of post-collisional Eocene UPVs in eastern

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32

771 Tibet (Campbell et al., 2014; Stepanov et al., 2014). Stepanov et al. (2017) suggested this

772 model could be applied to the Miocene trachyte suite of Tibet. The model requires two

773 principal stages: melting of blocks of continental crust within the mantle and variable

774 reaction between the rising crustal melts and the adjacent peridotitic mantle. Thus, felsic, K-

775 rich melts derived from continental crust thrust into the mantle interact with the overlying

776 mantle wedge during ascent to become mafic in composition (Stepanov et al., 2017).

777 However, several lines of evidence from the trachytic suites of southern Tibet preclude

778 this hypothesis (Wang et al., 2017a). For example, their εNdi values increase with increasing

779 SiO2 and do not trend toward Indian metasedimentary crust, as represented isotopically by

780 Himalayan leucogranites (Fig. 7C). Rather, the positive slope of SiO2 with Nd(i) in the

781 trachyte suites from southern Tibet (Fig. 7D) is more likely to be part of an array between

782 mafic magmas derived from an isotopically evolved mantle source, mixing with felsic

783 magmas derived from the isotopically unevolved southern Tibetan (Gangdese arc) crust. This

784 is consistent with the xenolith evidence from the trachytic suites (Wang et al., 2016).

785 Another critical point that conflicts with the Campbell et al. (2014) model is the variation

786 of trace elements on Harker diagrams. The trachytic suites show that the most mafic alkaline

787 rocks (with 10–12 wt. % MgO) have the highest concentrations of incompatible trace

788 elements (Figs. 2A-C, 4F, 8A). This incompatible enrichment in mafic trachytic rocks exists

789 across the entire spectrum of incompatible elements (Wang et al., 2017a, their Fig. 3B), and

790 shows that an enriched mantle source, not continental crust, controlled the primary trace

791 element geochemistry of the alkaline suite.

792 The consistently old Nd model ages (1.2–2.5 Ga) and relatively low zircon O isotopic

793 compositions (18O = 5–8.4‰; Wang et al., 2015a) for the trachytic suite, especially for the

794 more primitive variants (Fig. 7), suggest that these magmas were derived from low-degree

795 melting of an ancient (Proterozoic or Archean) SCLM. Furthermore, the SCLM must have

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33

796 undergone melt infiltration during the Proterozoic to explain the high HFSE and LILE

797 contents of the parental high-MgO but isotopically evolved magmas (sample T2A/98: MgO =

798 11.78 wt. %, εNdi = -18.5; Table A1). Melt infiltration, rather than hydrothermal fluid

799 metasomatism, is necessary because only melts can carry significant amounts of HFSE at

800 high P-T conditions (Spandler and Pirard, 2013). This melt infiltration may have occurred

801 during Proterozoic subduction-related magmatism, and is consistent with the typical

802 subduction-modified geochemical pattern of the trachytic rocks (Fig. 3D).

803 We propose therefore that the trachytic rocks originated from low degree partial melting

804 of SCLM at differing mantle depths. Two groups of trachytic suites can be distinguished on

805 La/Yb vs MgO and Sr/Y vs MgO plots (Fig. 8A and B). The eastern trachytic suite and

806 approximately half of the western suite plot on a very steep trend, with La/Yb ratios varying

807 between 30–200 for MgO <4 wt.%. By contrast, the other half of the suite, including the

808 transitional monzonitic types, have La/Yb ratios varying between 30–100 over an extended

809 MgO range (2–12 wt. % MgO). This suggests that the high-La/Yb trachytic magmas formed

810 in the garnet stability field, requiring depths of melting >70 km (Robinson and Wood, 1998),

811 whereas the low-La/Yb group probably formed at shallower mantle depths. This is consistent

812 with the presence of olivine in some rocks, suggesting the low-pressure melting reaction:

813 pyroxene + phlogopite = olivine + melt. The Sr/Y vs MgO plot (Fig. 8B) shows a similar

814 division into two groups. Overall, the contrasting La/Yb and Sr/Y ratios suggest that the

815 trachytic suites formed in a SCLM source region near the spinel–garnet transition.

816 The spinel-garnet transition implies mantle melting occurred between 60–80 km depth

817 (~2 GPa; Kinzler, 1997; Klemme and O’Neill, 2000). The geotherm for Tibet during the

818 Miocene, derived from xenoliths in the trachytic magmas, is ~16°C/km (Chen et al., 2009;

819 Wang et al., 2016), which suggests mantle temperatures at the Moho (~70 km depth) were

820 ~1100–1150°C. Recent melting experiments on phlogopite-bearing lherzolites and

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34

821 harzburgites (Condamine et al., 2016) showed that partial melting will occur in this range,

822 beginning at ~1000°C for 1 GPa, or at ~1150°C for 3 GPa, depending on fluorine content.

823 The K2O content of low-degree melts from phlogopite-lherzolite and phlogopite-harzburgite

824 is buffered between 6–8 wt.% (Condamine et al., 2016) similar to the values for the

825 trachyandesites (Fig. 2B). Also, the high K2O/Na2O ratios (2–8) of the trachytic suites,

826 typical of ultrapotassic rocks (Fig. 2D), are usually formed at low degrees of melting

827 (Condamine et al., 2016). The ~2 GPa (70 km) estimate for mantle melting suggested above,

828 combined with the relatively low inferred melting temperatures (~1100oC) and high

829 K2O/Na2O ratios, is in reasonable agreement with melting experiments on phlogopite-bearing

830 peridotites. Accordingly, the trachytic magmas are interpreted to have been produced by

831 partial melting of phlogopite-bearing (enriched) harzburgitic SCLM, which originally formed

832 in a supra-subduction environment during the Proterozoic.

833

834 6.4 Tibetan or Indian lithospheric mantle melting as a source for trachytes?

835 Evidence for the derivation of trachytic melts from an ancient, metasomatized lithospheric

836 mantle can be interpreted in two different ways: either the source was Tibetan SCLM in the

837 fore-arc region of the Gangdese arc, or Indian SCLM. Several authors, including some of us

838 (e.g., Wang et al., 2014c; Wang et al., 2015a), have proposed that the trachyte suites were

839 derived from Tibetan SCLM (Ding et al., 2003; Yang et al., 2015, 2016; Lu et al., 2015).

840 Here, we suggest the Indian plate SCLM as an alternative source.

841 Metasomatized Tibetan SCLM is an obvious potential source for the trachytic magmas,

842 but there are some issues with this model. Seismic studies show no evidence for the presence

843 of SCLM beneath Tibet today (Nábělek et al., 2009), and the isotopic record does not indicate

844 the participation of an older, evolved continental crust or underlying lithospheric mantle

845 throughout the magmatic history of the Gangdese arc, from ~200 Ma. Instead, the positive Hf

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846 isotopic composition of Gangdese arc magmas from Jurassic to early Eocene show repeated

847 reworking of juvenile crust with no ancient SCLM involvement (see Fig. 10 of Ji et al., 2009;

848 Fig. 4 of Liu et al., 2017).

849 A further point is that, if the trachyte suites originated from the Tibetan SCLM, it would

850 be difficult to keep producing these alkaline magmas over a ~20 m.y. period. As

851 demonstrated from experimental petrology (e.g., Wood and Turner, 2009; Condamine and

852 Médard, 2014; Condamine et al., 2016), generation of these magmas is by phlogopite

853 breakdown, producing low percentage melts derived from metasomatized SCLM (Foley,

854 1987), leaving a refractory harzburgitic residuum that cannot melt again under the moderate

855 16oC/km geothermal gradient constrained by the xenolith evidence.

856 From this we conclude that there was no old SCLM beneath the arc itself during its growth.

857 We argue instead that, if the evolved isotopic signal in the trachytes was derived from

858 Tibetan SCLM, it must have originally been part of the fore-arc region. We suggest that

859 during continental collision, the fore-arc lithosphere may have been thrust under the arc and

860 smeared northwards (Fig. 12c). Subsequent Miocene magmatism could then have involved

861 melting of this Tibetan SCLM and overlying Gangdese lower crust.

862 An alternative model, that the trachytes were derived from the Indian SCLM, is suggested

863 by the 20 m.y.-period of trachyte generation, which seems to require a continuously

864 rejuvenated source. We suggest that the subducting metasomatized Indian plate provides such

865 a source. Seismic and numerical models suggest that the Indian plate middle and upper crust

866 were mostly scraped off prior to subduction to form the Greater Himalaya accretionary prism

867 (e.g., Nábělek et al., 2009; Capitanio et al., 2010), leaving only the underlying mantle

868 lithosphere and part of lower crust as the main subducting component. The Indian plate

869 SCLM is thought to have undergone subduction metasomatism during the Proterozoic (Miller

870 et al., 2000), which led to its evolved Nd and Sr isotopic signature and K-rich, phlogopitic

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871 character (France-Lanord et al., 1988; Inger and Harris, 1993). Low-degree partial melting of

872 this material could have produced trachytic melts during subduction.

873 At this time we cannot distinguish between these two possible sources of Miocene

874 trachytic magmatism in Tibet (underthrust for-arc Tibetan SCLM or Indian plate SCLM).

875

876 6.5 Fluid-fluxed melting and oxidation of Tibetan lower crust in the Miocene: the origin

877 of high-Sr/Y granitoids

878 In this section, we demonstrate the importance of water in generating the Miocene high-

879 Sr/Y granitoids, and that the source was the Gangdese arc root. Zircons from the Miocene

880 high-Sr/Y granitoids provide temperatures that are generally <750°C (Fig. 6D, E). Maximum

881 zircon saturation temperatures (Tzr) for the eastern Miocene high-Sr/Y granitoids are also

882 <750°C (Fig. 9B), and the western high-Sr/Y group is marginally higher, up to 770°C

883 (although this difference is well within expected error for the method).

884 A requirement for accurate Tzr estimates is that the magma must be saturated in zircon

885 (Watson & Harrison, 1983). Calc-alkaline magmas generally reach zircon saturation when

886 SiO2 reaches values above 65 wt.% (Collins et al., 2016), which is reflected by a systematic

887 decrease in Zr content with silica increase in both the western and eastern high-Sr/Y granites,

888 from ~63–65 wt.% SiO2 to higher silica values (Fig. 9A). Given the presence of inherited

889 Paleogene zircons in high-Sr/Y granitoids (Wang et al., 2014a, Li et al. 2014), Tzr marks an

890 upper bound for temperature. A comparison of Ti-in-zircon temperatures with Tzr estimates

891 (Fig. 9C) shows they yield similarly low-T values. The same does not hold for the Paleocene-

892 Eocene Gangdese arc rocks, which show poor agreement between Ti-in-zircon and Tzr

893 temperature estimates in Figure 8C. However, many of these samples were undersaturated in

894 zircon, as shown by their increasing Zr content up to ~65 wt.% SiO2. The comparisons

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895 between Tzr and Ti-in-zircon temperature estimates and silica contents reinforces the point

896 that the high-Sr/Y granitoids were formed from cool magmas.

897 Given that dehydration melting of the crust requires temperatures of at least 850°C to

898 generate reasonable granitic magma volumes, the consistently low temperatures for the

899 Miocene high-Sr/Y granitoids and transitional monzonites (Fig. 9B,C), requires additional

900 water-fluxed melting (cf. Weinberg and Hasalová, 2015; Collins et al., 2016). Plagioclase

901 compositions in the eastern high-Sr/Y granitoids further support this interpretation because

902 they show excess aluminium (Fig. 10), which has been linked to high melt water contents

903 (Williamson et al., 2016).

904 Evidence from the Paleocene–Eocene Gangdese arc rocks suggests that the Tibetan arc

905 root was relatively reduced, but became moderately oxidized and strongly hydrated during

906 Oligocene–Miocene flat continental subduction. This lithospheric metasomatism is thought to

907 have played an important role in the subsequent magmatic flare-up and associated porphyry

908 copper mineralization in the Miocene (Wang et al., 2014b). The oxidation process can be

909 dated back to the start of Indian plate subduction in the Eocene, when a range of sediments,

910 including evaporites and carbonates that might have existed on the leading edge of the Indian

911 passive margin were underthrust beneath Tibet. These variably oxidized metasedimentary

912 rocks had the potential to oxidize the arc root as the underthrust Indian plate progressively

913 dehydrated during continental collision (e.g., Massonne, 2016).

914

915 6.6 Mixing model for Miocene high-Sr/Y magmas

916 Whole-rock and zircon isotopic compositions for Miocene high-Sr/Y granitoid samples

917 are generally similar to those of the Paleocene-Eocene Gangdese arc (Wang et al., 2015a),

918 suggesting derivation by partial melting of the arc root (Hou et al., 2004, 2015). The

919 distinctively higher La/Yb of the eastern granitoids compared to the Gangdese arc rocks (Fig.

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38

920 7B, 9A) accords with their distinctively high-Sr/Y contents (Fig. 4A, B, 7A), but many other

921 major and trace elements overlap with the felsic endmembers of the Gangdese arc (Figs. 2, 3,

922 4, 8). The high La/Yb and Sr/Y ratios demonstrate that these Miocene granitoids formed by

923 melting of the Paleocene-Eocene arc, within the garnet stability field at depths >30 km.

924 In detail, however, there are some subtle but significant differences between the high-Sr/Y

925 granitoids in eastern and western Tibet, best revealed by the variation of εNdi values (Figs. 5,

926 7). Although the eastern high-Sr/Y granitoids have values that overlap with those of the

927 Paleocene-Eocene Gangdese arc (+8 to -5), the range extends beyond the Gangdese envelope

928 toward the trachyte suites. This is also evident geographically, with the high-Sr/Y granitoids

929 being most primitive in the east, and becoming progressively more evolved toward the west

930 (Fig. 5B, C). Chemically and geographically, the trend is toward the evolved isotopic

931 compositions of the trachytic suite, indicating that some high-Sr/Y granitoids have

932 incorporated trachytic components.

933 The Sr/Y vs Ndi plot (Fig. 7A) most convincingly shows the chemical-isotopic

934 interrelationship. High-Sr/Y Miocene granitoids are chemically distinctive from the

935 Gangdese arc rocks, but isotopically similar. On the other hand, the trachytic suite varies

936 between low- and high-Sr/Y at evolved Nd isotopic compositions, with a general trend

937 toward the eastern high-Sr/Y Miocene granitoids. This suggests that variable degrees of

938 mixing occurred between low-Sr/Y evolved trachyandesites, and high-Sr/Y crustal melts

939 derived from Gangdese crust.

940 Isotopic variations in the eastern Miocene high-Sr/Y granitoids are mimicked by other

941 compositional variations. A clear demonstration is the compositional overlap of MgO and

942 SiO2 between the felsic members of the trachytic suites and the mafic endmembers of the

943 eastern Miocene high-Sr/Y granitoids (Fig. 7C, D). The variation is also evident for trace

944 elements such as Cr (Fig. 7E) and Th (Fig. 7F), and with Zr (Fig. 9A): Zr drops from ~1000

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945 ppm to 100 ppm as SiO2 increases from 55–70 wt. %. The order of magnitude drop in Zr

946 content corresponds to a change from εNd of -15 to +2 (Fig. 9D), demonstrating the extreme

947 effect that mixing had on trace element abundance, rather than crystal fractionation.

948 Intermediate between the mantle-derived (trachytic) endmember and the crustally derived

949 most felsic, silica-rich endmembers of the high-Sr/Y granitoid suite, lie the western Miocene,

950 high-Sr/Y granitoids and the transitional monzonite suite. These granitoids tend to overlap

951 with the lower silica, lower alkali group of the eastern granitoids (Fig. 2), also evident in their

952 slightly higher CaO, MgO, Ni (Fig. 4) and Zr (Fig. 9D) contents. Considering the two

953 hypothetical endmembers defined in Fig. 5A, most of the eastern Miocene granitoids contain

954 5–10 % trachytic magmatic component, whereas most of the western Miocene granitoids

955 contain 15–20 % of that component. Overall, the Miocene high-Sr/Y magmas are dominated

956 by a Gangdese arc source, with 5–20 % contamination by mantle-derived, alkaline magmas

957 of the coeval trachytic suites.

958 An antithetic geographical relationship exists between Miocene trachytic and high-Sr/Y

959 granitoid suites in southern Tibet. Whereas the high-Sr/Y granitoids are voluminous in the

960 east, they are sparse in the west. Conversely, trachytic rocks are sparse in the east, but much

961 more voluminous in the west. The isotopic array indicates the two endmembers mixed more

962 effectively in the east (Fig. 5B, C), producing the greater range in isotopic compositions of

963 the eastern high-Sr/Y granitoids, and the limited number of erupted alkaline magmas were

964 much more thoroughly mixed than those in the west. No eastern trachytic magmas have εNd

965 values <-9, whereas the isotopic range extends to -18 farther west, in regions where high-

966 Sr/Y plutons are not present. The antithetic relationship suggests that the more voluminous

967 high-Sr/Y crustal melts in the east acted as a rheological and probably density barrier to

968 ascending mantle-derived melts.

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969 In summary, there is evidence for variable degrees of hybridization between mantle-

970 derived trachytic magmas and high-Sr/Y granitoid melts formed by anatexis of the older

971 Gangdese magmatic arc in a water-fluxed environment (Wang et al., 2016, 2017a). This

972 model is most similar to that of Yang et al. (2015, 2016), and contrasts with many other

973 petrogenetic models (Qu et al., 2004, 2007; Gao et al., 2007, 2010; Chung et al., 2003; Hou et

974 al., 2004; Guo et al., 2007; Hou et al., 2009; Li et al., 2011; Wang et al., 2014a, b; Zheng et

975 al., 2012; Liu et al., 2017). Yang et al.’s (2015) model was based on the chemical features of

976 the transitional Qulong “high Mg-diorite” (our transitional monzonite group) (Fig. 1). It

977 differs from the model presented here in that we suggest the mantle source was the Indian

978 SCLM or Tibetan fore-arc SCLM. In addition, our model differs from Yang and co-workers

979 in that we suggest the steepening of the subducting Indian plate in the early Miocene allowed

980 temperatures to rise in the lower Tibetan crust as an asthenospheric mantle wedge began

981 to open, which induced voluminous deep crustal melting in the eastern Gangdese. Trachytic

982 magmas would have transported additional heat and fluid into these lower crustal melting

983 zones, and may have enhanced the melting process.

984 A mixing model has also been proposed by Liu et al. (2017) for the Miocene magmatic

985 rocks of southern Tibet, but their model contrasts with that presented herein, because they

986 consider that the Miocene high-Sr/Y granitoids were derived from the Indian plate and the

987 potassic volcanic rocks from Tibetan crust. Both schools of thought agree the “ultrapotassic”

988 component of the trachytic alkaline magmas (see subdivision between potassic and

989 ultrapotassic in Fig. 2D) were mantle derived, but Liu et al. (2017) discriminate the “potassic

990 rocks” from the “ultrapotassic group” (our trachytic suite) and suggest the former ultimately

991 had a crustal origin. The authors argued that “poorly varying” (presumably meaning a narrow

992 range of values) and negative zircon εHf(t) values of the potassic volcanic rocks provided

993 clear evidence for derivation from ancient Lhasa terrane crust with minor input from related

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994 ultrapotassic magma. However, as shown in Fig. 7, both the Indian crust (represented by the

995 Himalayan leucogranites) and the Indian SCLM (represented by the 3.35 Ga komatiites,

996 Jayananda et al., 2008), have highly negative εNd(t) values. Thus, although Liu et al. (2017)

997 assume that the dramatic decrease of co-magmatic zircon εHf(t) values since ~35 Ma is

998 strong evidence for the enhanced mass transfer from underthrusted Indian continental crust, it

999 is equally possible that it represents an influx of trachytic magma (their ultrapotassic magmas)

1000 from Indian or Tibetan ancient SCLM.

1001 Most significantly, the diagrams in Fig. 7 indicate that the silicic endmember involved in

1002 the process is close to juvenile in terms of its isotopic signature, similar to the Paleocene-

1003 Eocene Gangdese granitoids, whereas the most mafic, Mg-rich magmas were the most

1004 isotopically evolved. Thus, Indian crustal magmas with evolved signatures, such as the

1005 Himalayan leucogranites (Fig. 7A, C, D) could not be a significant endmember in the

1006 hybridization process that generated the high-Sr/Y granitoids, whereas granitic magmas

1007 derived from anatexis of the Paleocene-Eocene Gangdese arc could. Conversely, the trachyte

1008 suites, derived from melting of harzburgite of the Indian SCLM, are both Mg-rich and

1009 isotopically evolved and would explain both the compositional and isotopic variation in the

1010 Miocene high-Sr/Y granitoids.

1011 Liu et al. (2017) also used variation in the alkali ratio (K2O/Na2O) against Y and SiO2

1012 (their fig. 10) to suggest that the SCLM-derived trachytic (their ultrapotassic) melts only

1013 played a minor role in Miocene granitoid magmatism. They suggested that K2O/Na2O in the

1014 high-Sr/Y granitoids (their adakites) increases with increasing SiO2 and decreasing Y.

1015 However, this increase only applies to the low-K, Paleocene-Eocene Gangdese arc rocks, not

1016 to the Miocene granitoids (Fig. 11); the Miocene trachyte-granitoid array shows a steady

1017 decrease of K2O/Na2O with increasing silica and decreasing Y, as predicted by magma

1018 mixing models.

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1019 Liu et al. (2017) further suggest that the potassic volcanism is dominated by recycling of

1020 the Lhasa terrane crust because the negative and variable co-magmatic zircon Hf isotopic

1021 variations are comparable with detrital zircon records and magmatic zircons from the

1022 Mesozoic granitoids outcropping in the central and northern Lhasa subterranes (cf., Liu et al.,

1023 2014). We concur with Liu et al. (2014) that such zircons are mostly crust-derived (high

1024 U/Yb) xenocrysts entrained within mantle-derived ultrapotassic (trachytic) magmas and that

1025 their heterogeneous Hf isotopes indicate assimilation of Lhasa terrane crust during ascent of

1026 those magmas. Indeed, the highlighted granitic xenoliths (Liu et al., 2014, their Fig. 1) are

1027 non-foliated, upper crustal fragments, indicating late entrainment of Lhasa terrane within the

1028 crustal column rather than derivation from lower crustal sources. The presence of these upper

1029 crustal xenoliths and xenocrysts in the trachytic suites does not indicate that the potassic (or

1030 ultrapotassic) magmas were derived from Tibetan crust.

10311032 6.7 Thermal structure of the Miocene Gangdese belt

1033 Mantle xenoliths entrained in the Miocene trachytic magmas place constraints on

1034 petrological and geophysical models for that time interval in southern Tibet. Mafic granulite

1035 xenoliths from within the trachytic suites in the western Gangdese belt yielded temperatures

1036 of 1130–1330o C and pressures between 22 and 26 kbar, defining a geotherm of ~16oC km-1,

1037 suggesting that mafic crust extended to between 70–85 km depth beneath Tibet during the

1038 Miocene (Wang et al., 2016, 2017a). Hydrous ultramafic xenoliths have abundant hornblende

1039 and contain ~85 Ma-old zircons typical of the Gangdese arc (Chan et al., 2009), likely

1040 representing a deep cumulate section of the arc.

1041 Felsic granulite xenoliths indicate that the Miocene Tibetan arc root had a basal temperature

1042 of ~850°C (Wang et al., 2016), close to biotite dehydration melting temperatures. Assuming a

1043 50 km thick Tibetan crust, as indicated by seismic data for the present (Nábelek et al., 2009),

1044 the felsic granulite xenoliths yield a geothermal gradient of 17°C /km. We use this gradient to

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1045 constrain the 750°C isotherm at shallower crustal levels (Fig. 12), which is the approximate

1046 temperature of the high-Sr/Y granitoid magmas estimated above. Farther north, deep seismic

1047 experiments beneath the Qiangtang and Songpan-Ganzi terranes identified the lithosphere-

1048 asthenosphere boundary (LAB) at ~80 km depth (Owens and Zandt, 1997). Assuming the

1049 LAB (lithosphere-asthenosphere boundary) is a thermal boundary layer at ~1100°C, the

1050 geothermal gradient is ~16°C/km, similar to the other estimated gradients. We use this to

1051 locate the LAB beneath central Tibet.

1052 Another thermal constraint for Fig. 12 comes from experimental melting models for

1053 phlogopite-bearing peridotites. Wendlandt and Eggler (1980) showed that the beginning of

1054 melting of a phlogopite-bearing spinel lherzolite under anhydrous conditions was at ~1075 oC

1055 at 10 kbar and ~1120 oC at 20 kbar, suggesting that under Tibet at ~70 km depth (~20 kbar),

1056 the temperature was approximately 1100 °C in agreement with the 16 °C geothermal gradient

1057 determined from xenoliths within the trachytic magmas. This result provides confidence for

1058 the inferred thermal structure beneath Tibet, and is also consistent with the

1059 geochemical/petrological arguments that trachyte generation occurred near the garnet-spinel

1060 transition, at ~70 km.

1061 The thermal structure constrained by petrological arguments appears to be unlike the

1062 structure given in the classical papers by Toksöz et al. (1971), Bird et al. (1975), and Peacock

1063 (1990), and more recently by Beaumont and co-workers (Warren et al., 2008; Beaumont et al.,

1064 2009). All these models project a cold slab extending deep into the mantle, where

1065 temperatures can remain at ~600oC to at least 100 km depth, but also suggest the lithospheric

1066 mantle wedge is much less than 1000oC, for any condition involving continental collision.

1067 Therefore, such models cannot predict the generation of hot, trachytic magmas beneath

1068 southern Tibet. If the trachytic magmas are derived from the slab (our hypothesis), then the

1069 slab interior heats up much more rapidly than predicted by the thermal models (Fig. 12). In

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1070 the next section, we use these petrological constraints to derive a geodynamic model for Tibet

1071 between 15–20 Ma.

1072

1073 6.8 Geodynamic model

1074 Receiver-function images (Kumar et al., 2006; Zhao et al., 2010) and body and surface

1075 wave tomographic models (Nunn et al., 2014), suggest a west to east increase in the angle of

1076 dip of the Indian plate lithosphere, and a decrease in thickness of the Indian plate lithosphere,

1077 both west and east (Fig. A1). The flat subduction inferred for the early stages of collision

1078 (between ~50 and 30 Ma, section 3.2) differs from present-day geometry, particularly for the

1079 east. The geometries of the subducting Indian plate today (Fig. A1) are equivalent to type IIa

1080 “continental subduction” in the east, versus type IIb “continental underthrusting” in the west

1081 (Massonne, 2016, his Fig. 1). We suggest that the steepening of the Indian plate in eastern

1082 Tibet occurred in the middle Miocene, and was related to the generation of voluminous high-

1083 Sr/Y granitoid magmatism in that region. A coherent geodynamic model must not only take

1084 into account the contrasting geometry of the subducting Indian plate from west to east, but

1085 also the increased volumes of high-Sr/Y granitoids and the greater numbers of PCDs in the

1086 east, and the diminished volumes of trachytic rocks.

1087 The following scenario is proposed:

1088 1. During the Paleocene–Eocene (pre-50 Ma), normal, subduction-related Gangdese arc

1089 magmatism occurred. The medium- to high-K calc-alkaline arc magmas (Fig. 2) were

1090 relatively reduced (Fig. 6). Starting at ~50 Ma, the Eocene magmas are characterized

1091 by higher Th/Y and La/Yb ratios, suggesting crustal thickening.

1092 2. The slab began to flatten during the Oligocene as Indian lithosphere entered the

1093 subduction zone and a period of magmatic quiescence ensued. Crustal thickening

1094 occurred and the Himalayan orogenesis began. Although some metasedimentary

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1095 materials reach the lower crust based on the high δ18O values (up to 8.03) of olivine

1096 from xenoliths hosted by trachytes (Liu C et al., 2014), most of the upper crust of the

1097 Indian plate was scraped off (Capitanio et al., 2010) and formed the accretionary

1098 wedge of Tethyan metasedimentary rocks south of the Indus-Yarlung Tsangpo suture

1099 zone. Strong coupling between the two plates during this flat subduction event

1100 removed ancient Tibetan SCLM beneath southern Tibet, consistent with the seismic

1101 evidence (Nábĕlek et al., 2009).

1102 3. During this flat subduction mode, the geothermal gradient in the Tibetan crust

1103 decreased from typically 30–40oC/km or greater during Gangdese arc magmatism to

1104 ~16oC/km by the Miocene during continental underthrusting, based on xenoliths from

1105 the western Gangdese belt (Chan et al., 2009; Wang et al., 2016).

1106 4. During ongoing continental underthrusting, the remnant subducted Indian lower crust

1107 and some retained metasediments from upper crust progressively dehydrated as it

1108 converted to eclogite facies. Massonne (2016) demonstrated that sedimentary rocks in

1109 a subducting slab can release up to 2.5 wt. % water at ~600 °C over a wide range of

1110 pressures, with the aqueous fluids rising into and hydrating the overlying rocks.

1111 Serpentinised peridotites also released significant water volumes during serpentine

1112 breakdown, which occurs at <650°C for P < 2.0 GPa. At ~50 km depth, the mantle is

1113 saturated at 1.8 wt. % H2O at 600 °C (Massonne et al., 2016), allowing a free aqueous

1114 phase to rise into the overlying lithosphere.

1115 5. The northern margin of the Indian continent was covered by carbonates and evaporitic

1116 sediments (Mukherjee et al., 2003; Scheibner and Speijer, 2008; Johnston et al.,

1117 2011). They are also the most ductile of supracrustal rocks, and commonly define

1118 fault structures. Although most of upper crust has been scrapped off, some of these

1119 sediments can persist to great depths. During Indian flat subduction, dehydration of

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1120 these carbonates and evaporites (and other sediments) led to further hydration and

1121 oxidation of the base of the Gangdese arc. This oxidation step is critical to subsequent

1122 metallogeny, because the Gangdese arc lower crust is thought to have been relatively

1123 reduced prior to this time, and did not generate magmas that were fertile for porphyry

1124 formation.

1125 6. Steepening of the Indian subduction in the east (Fig. 12B, C) resulted in the opening

1126 of an asthenospheric mantle wedge, and caused temperatures to rise in the overlying

1127 lower Tibetan crust (compare Fig. 12A, B, C). This induced melting of water-fluxed

1128 (or hydrated) lower crust at the ambient temperature of ~800°C at depths of 45–50

1129 km.

1130 7. Early Miocene trachytic magmatism began along the southern margin of Tibet as a

1131 result of phlogopite breakdown at ~1100°C, near the garnet-spinel transition (~70

1132 km), producing low volume, K-rich, trachytic partial melts (Condamine et al., 2016).

1133 The location of the trachytic magmas in southern Tibet suggests that the ~1100°C

1134 isotherm was at shallower depths than estimated from thermo-mechanical models of

1135 subduction (Fig. 12).

1136 8. The highly evolved Nd and Sr isotopic signature of the trachytic magmas, yielding

1137 Early Proterozoic TDM model ages (Fig. 5), and depletions for Nb and Ta in mantle-

1138 normalized trace element patterns (Fig. 3), indicate that the SCLM was originally

1139 metasomatised during Proterozoic suprasubduction zone magmatism. As mentioned

1140 in section 6.4, there are two possible sources for trachytic melts: ancient Indian

1141 SCLM, or ancient Tibetan SCLM in the fore-arc region.

1142 9. The impact of hot, rising trachytic magmas into Tibetan crust differs from west to

1143 east. In the west, the melts rose into a cool, relatively rigid lower crust still cool

1144 because of ongoing flat slab subduction (Fig. 12A). Accordingly, extensive crustal

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1145 melting did not occur and trachytic melts could continue toward the surface virtually

1146 unmodified. This explains the relatively high proportion of trachytic intrusions

1147 relative to high-Sr/Y granitoids in western Tibet.

1148 10. By contrast, the resultant crustal melting was extensive caused trapping of trachytic

1149 melts, resulting in their sparse eruption at surface, but evidence for mixing with the

1150 granitoid melts. Fluids released from these trachytes may have enhanced fluid-fluxed

1151 melting of the metasomatized and now warm Tibetan lower crust.

1152 11. Trachytic magmas were not the oxidizing agent of the Gangdese lower crust, because

1153 the dominant (western) trachytic suite rocks are significantly more reduced than the

1154 high-Sr/Y granitoids (Fig. 6C). The western trachytic and transitional monzonitic

1155 magma have much higher Ce4+/Ce3+ ratios but broadly similar EuN/EuN* ratios with

1156 the western trachytic magmas, suggesting that neither of them was significantly

1157 oxidized. Instead, we propose that the oxidation evident in the high-Sr/Y granitoids

1158 occurred during metasomatism of the Tibetan lower crust by fluids released from the

1159 Indian plate during flat subduction.

1160 12. Ascent of trachytic magma into the lower Tibetan crust was accompanied by high

1161 degrees of olivine fractionation until the magmas reached low-MgO (~2 wt. %) and

1162 high silica (~65–70 wt. %) contents (Fig. 7C, D). At this stage, the magmas were able

1163 to release water to stimulate crustal melting, and at the same time were capable of

1164 mixing with lower crustal melts (Figs. 5, 7) at ~750–800oC, producing some high-

1165 Sr/Y granitoid magmas with high Cr-Ni-Mg# contents (Figs. 2D, 7F, 8D).

1166 13. The high-Sr/Y granitoid magmas derived from partial melting of hydrated and

1167 oxidized Gangdese arc base were capable of scavenging Cu from originally sulphide-

1168 rich, probably metalliferous portions of the Gangdese arc cumulates in the lower crust

1169 and/or lithospheric mantle. The metals were remobilized under these relatively

27152716271727182719272027212722272327242725272627272728272927302731273227332734273527362737273827392740274127422743274427452746274727482749275027512752275327542755275627572758275927602761276227632764276527662767276827692770277127722773

48

1170 oxidised melting conditions, and transported into the upper crust by the high-Sr/Y

1171 granitoid magmas to become the key components of the PCDs in eastern Tibet. The

1172 more extensive crustal melting in the eastern Gangdese (compared to the west)

1173 explains the occurrence of PCDs

1174

1175 6.9 Metallogenic implications

1176 Similar to PCDs in arc settings, Gangdese post-collisional PCDs are also associated with

1177 hydrous and oxidized magmas (Hou et al., 2015; Lu et al., 2015; Yang et al., 2015; Wang et

1178 al., 2015a). Although partial melting of subduction-modified lower crust has been proposed

1179 to generate such magmas, recent studies (Lu et al., 2015; Yang et al., 2015) question the

1180 ability of dehydration melting of garnet amphibolite in a thickened lower crust to generate

1181 sufficient quantities of hydrous magma to form porphyry deposits upon upper crustal

1182 emplacement. In contrast with Yang et al. (2015), we suggest the shallow-subduction

1183 metasomatism was not enough to cause melting until slab steepening in the Miocene caused

1184 heating. However, the shallow subduction was a very important precursor event, which

1185 rendered the lower crust fusible (and oxidized) when temperatures rose as the asthenospheric

1186 mantle wedge opened. This process of hydrating the Gangdese arc base can be dated back to

1187 the Eocene, when the Indian plate flat subduction started, and lasted till the Miocene. We

1188 have also suggested above that mixing with trachytic magmas is necessary to explain some

1189 Miocene granitoid magmas with enriched Cr and Ni contents, and high Mg# at 56–70 wt. %

1190 SiO2. The arrival of trachytic magmas at the base of the Tibetan arc was broadly coincident

1191 with the melting of the hydrated base of arc. Trachytic magmas may have been at least in part

1192 the heat and fluid source that triggered or enhanced melting at the base of the arc.

1193 A key question is how did Miocene high-Sr/Y granitoid magmas, derived from the melting

1194 of reduced cumulates of the Gangdese arc root (ΔFMQ -1.2 to +0.8; Wang et al., 2014c)

27742775277627772778277927802781278227832784278527862787278827892790279127922793279427952796279727982799280028012802280328042805280628072808280928102811281228132814281528162817281828192820282128222823282428252826282728282829283028312832

49

1195 become oxidized? As discussed in sections 6.2, 6.5 and 6.8, we consider that subduction of a

1196 range of Tethyan sediments from the Indian plate oxidized the roots of the arc ahead of

1197 Miocene melting.

1198 Miocene high-Sr/Y magmas are oxidized and can carry more sulphur than their reduced

1199 counterparts (Wang et al., 2014b; Hou et al., 2015, Tomkins et al., 2012), and could have

1200 scavenged sulphides and their metals that became trapped within the roots of the reduced

1201 Gangdese arc. The assimilation of sulphides would have limited reducing effect on the high-

1202 Sr/Y magmas (Tomkins et al., 2012), which were capable of transporting metals to give rise

1203 to PCDs (Richards, 2011; Chiaradia et al., 2012).

1204 The observation that PCDs are restricted to the eastern section of the Gangdese belt is a

1205 direct result of the increased crustal melting in the east as the asthenospheric mantle wedge

1206 opened during the transition from flat to steep subduction. In contrast, the lack of

1207 asthenospheric mantle wedge in the west restricted the degree of crustal melting and limited

1208 the ability to trap trachytic melts, leading to the generation of few poorly mineralized

1209 granitoids but the eruption of large volumes of trachytic volcanic rocks at surface.

1210

1211 Acknowledgements

1212 This research was financially cosupported by the MOST of China (2016YFC0600304 and 20

1213 16YFC0600407), the Chinese National Natural Science Foundation (91755207 and 4122500),

1214 and the 111 Project (B18048). William J. Collins was supported by ARC grant DP

1215 120104004, and Jeremy P. Richards was supported by Discovery Grant from the Natural

1216 Sciences and Engineering Research Council of Canada. Wen-yan He and Qiuyun Li are

1217 thanked for zircon trace element analysis, and Andrew Locock is thanked for assisting EPMA

1218 analysis. Andreas Audétat is thanked for reviewing an early version of the manuscript.

1219

28332834283528362837283828392840284128422843284428452846284728482849285028512852285328542855285628572858285928602861286228632864286528662867286828692870287128722873287428752876287728782879288028812882288328842885288628872888288928902891

50

1220

1221 Figures Captions and Tables

1222

1223 Figure 1. Geology of the Gangdese magmatic belt in the Lhasa terrane, showing the

1224 distribution of Paleocene-Eocene Gangdese magmatism (including intrusions and Linzizong

1225 volcanic successions), Miocene high-Sr/Y granitoids, and western (W) Miocene trachytic

1226 rocks, and eastern (E) Miocene trachydacites (sparse). Notice how Miocene granitoids west

1227 of ~88°E are sparse and mostly barren, whereas in the east they are more abundant and

1228 commonly associated with PCDs. Blank areas are Quaternary successions. Map modified

1229 from: Hou et al. (2004); Zhao et al. (2009); Wang et al. (2015a).

1230

1231 Figure 2. (A) Total alkali-silica (TAS) diagram, (B) K2O vs SiO2 plot, (C) K2O vs MgO plot

1232 and (D) K2O vs Na2O plot for the six main Cenozoic suites in the Gangdese belt. References

1233 for data provided in Table A1.

1234

1235 Figure 3. (A-C) Chondrite-, and (B-D) N-MORB-normalized trace element diagrams for the

1236 six main Cenozoic suites in the Gangdese belt. Normalization values are from Sun and

1237 McDonough (1989). References for data provided in Table A1.

1238

1239 Figure 4. Major and trace element plots showing the features of the six main Cenozoic suites

1240 in the Gangdese belt: (A) Sr/Y vs Y, (B) Sr/Y vs La/Yb, (C) Mg# vs SiO2, (D) CaO vs Al2O3,

1241 (E) MgO vs SiO2, and (F) Th vs Ni. References for data provided in Table A1.

1242

1243 Figure 5. (A) εNdi (T=15Ma) vs (87Sr/86Sr)i (T=15Ma). (B) εNdi (T=15Ma) vs. longitude and

1244 (C) NdTDM2 vs longitude for the six main Cenozoic suites in the Gangdese belt. Note: the

28922893289428952896289728982899290029012902290329042905290629072908290929102911291229132914291529162917291829192920292129222923292429252926292729282929293029312932293329342935293629372938293929402941294229432944294529462947294829492950

51

1245 Paleocene-Eocene Gangdese arc calc-alkaline values were calculated at 15 Ma for

1246 comparison to Miocene rocks. In A, grey line shows mixing between a primitive end-member

1247 for the arc root, and a primitive trachyte for melts derived from Indian lithospheric mantle or

1248 ancient fore-arc Tibetan SCLM. The values for end-members used in this mixing model are:

1249 (87Sr/86Sr)i = 0.703 and εNdi = +8.5 (the most primitive Gangdese Paleocene-Eocene

1250 Gangdese arc sample), and average Sr (623 ppm) and Nd (20.3 ppm) values from eastern

1251 Miocene high-Sr/Y granitoids. Values for the primitive trachytic melts are taken from the

1252 western Miocene suite: (87Sr/86Sr)i = 0.726 and εNdi = -17, average Sr (916 ppm) and Nd

1253 (137 ppm). References for data provided in Table A1.

1254

1255 Figure 6. Trace element composition of zircons from the six main Cenozoic suites in the

1256 Gangdese belt: (A) Chondrite-normalized REE diagram; inset is a plot of Dy/Yb vs. Yb

1257 indicating that the western trachytes have low Yb contents and a Dy/Yb ratio close to unity,

1258 contrasting with all other suites; (B) Ce4+/Ce3+ histogram showing all individual analyses; (C)

1259 Ce4+/Ce3+ vs. Eu/Eu*; (D) Th/U vs T(ti-zr) (Ti-in-zircon temperature based on equations

1260 from Watson and Harrison, 2005); (E) Ce4+/Ce3+ vs. T(ti-zr); (F) Ce4+/Ce3+ vs Th/U.

1261 Normalization values are from Sun and McDonough (1989). Chemical values for zircons in

1262 C-F are averages of several spots in single sample from which averages and errors are

1263 calculated. References for data provided in Table A2. Averages of Qulong and Jiama ore-

1264 forming samples are from Lu et al. (2016).

1265

1266 Figure 7. Major and trace element plots for the six main Cenozoic suites in the Gangdese belt:

1267 (A) Sr/Y vs. εNdi (T=15Ma), (B) La/Yb vs. εNdi (T=15Ma), (C) SiO2 vs. εNdi (T=15Ma),

1268 (D) MgO vs. εNdi (T=15Ma), (E) Cr vs. εNdi (T=15Ma); (F) Th vs. εNdi (T=15Ma).

1269 References for data provided in Table A1.

29512952295329542955295629572958295929602961296229632964296529662967296829692970297129722973297429752976297729782979298029812982298329842985298629872988298929902991299229932994299529962997299829993000300130023003300430053006300730083009

52

1270

1271 Figure 8. Major and trace element plots for the six main Cenozoic suites in the Gangdese belt:

1272 (A) La/Yb vs. MgO and (B) Sr/Y vs MgO. References for data provided in Table A1.

1273

1274 Figure 9. Zr content and T estimates: (A) whole rock Zr vs SiO2 plot; (B) Tzr (Zr saturation

1275 temperature) vs SiO2; (C) T(ti-zr) (Ti-in-zircon temperature) vs Tzr; (D) whole rock Zr vs

1276 εNdi (T=15Ma) for the six main Cenozoic suites in the Gangdese belt. T(ti-zr) estimation

1277 based on equations from Watson and Harrison (2005), Tzr temperature estimate based on Zr

1278 saturation (Boehnke et al., 2013).

1279

1280

1281 Figure 10. Al/(Ca+Na+K) vs. An% for plagioclase from different suites. Plagioclase crystals

1282 from eastern Miocene high-Sr/Y rocks plot above the line, indicating they have excess Al

1283 (Williamson et al., 2016), unlike Paleocene-Eocene Gangdese arc and western trachyte

1284 plagioclase. References for data in plots provided in Table A3.

1285

1286 Figure 11. Major and trace element plots for the six main Cenozoic suites in the Gangdese

1287 belt: (A) K2O/Na2O vs. SiO2 and (B) K2O/Na2O vs. Y. References for data provided in Table

1288 A1.

1289

1290 Figure 12. Cartoon illustrating the contrasting subduction geometry associated with Miocene

1291 magmatism in the Gangdese belt: (A) Flat subduction and continental underthrusting in the

1292 western Gangdese belt, and (B-C) steeper subduction in the eastern Gangdese belt, consistent

1293 with deep seismic experiments. Both geometries led to underplating of oxidized supracrustal

1294 rocks, and generated a thermal structure capable of hydrating and oxidizing the overlying

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53

1295 Tibetan plate lithosphere. Upon opening of the asthenospheric mantle wedge in the east in the

1296 Miocene, the hydrated and oxidized Tibetan lower crust began to melt to form high-Sr/Y

1297 magmas. This process was accompanied by the formation of trachytic partial melts in the

1298 underthrust Indian plate SCLM (B) or for-arc Tibetan SCLM (C), which also invaded the

1299 Tibetan lower crust. In the east, these hydrous, alkaline magmas were trapped by the lower

1300 crustal melt sheets, where they released fluids that contributed to further fluid-fluxed crustal

1301 melting, and variably mixed with these crustal melts. The increased oxidation state of the

1302 Tibetan lower crust caused metals trapped in sulfides from previous subduction-related

1303 magmatism to be remobilized, generating magmas that were fertile for porphyry Cu deposit

1304 formation. In contrast, the lack of an asthenospheric mantle wedge in the west caused more

1305 limited crustal melting that failed to trap the ascending trachytic magmas. Consequently, only

1306 small volumes of crustal melt (high-Sr/Y granitoid) were generated, with only Zhunuo Cu-

1307 Mo deposit and a few intrusions with mineralization, but large volumes of trachytic volcanic

1308 rocks were erupted.

1309

1310 Table 1 Geochemical comparison of Gangdese Cenozoic igneous rocks and their zircons.

1311

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1678 Post-collision, shoshonitic volcanism on the Tibetan Plateau: implications for convective 1679 thinning of the lithosphere and the source of ocean island basalts. J. Petrol. 37, 45–71.1680 Tilmann, F., Ni, J., INDEPTH III Seismic Team, 2003. Seismic imaging of the downwelling 1681 Indian lithosphere beneath Central Tibet. Science 300, 1424–1427.1682 Tomkins, A.G., Weinberg, R.F., McFarlane, C.R.M., 2009. Preferential magma extraction 1683 from K- and metal-enriched source regions in the crust. Miner. Deposita 44, 171–181.1684 Van der Voo, R., Spakman, W., Bijwaard, H., 1999. Tethyan subducted slabs under India. 1685 Earth Planet. Sci. Lett. 171, 7–20.1686 van Hinsbergen, D.J.J., Lippert, C., Dupont-Nivet, G., McQuarrie, N., Doubrovine, V., 1687 Spakman, W., Torsvik, T.H., 2012. Greater India basin hypothesis and a two-stage 1688 Cenozoic collision between India and Asia. PNAS 20, 7659–7664.1689 Wagner, L.S., Beck, S., Zandt, G., 2005. Upper mantle structure in the south central Chilean 1690 subduction zone (30° to 36°S). J. Geophys. Res. 110, B01308.1691 Wang, R., Richards, J.P., Hou, Z.Q., Yang, Z.M., DuFrane, S.A., 2014a. Increased magmatic 1692 water content–the key to Oligo-Miocene porphyry Cu–Mo ± Au formation in the eastern 1693 Gangdese belt, Tibet. Econ. Geol. 109, 1315–1339.1694 Wang, R., Richards, J.P., Hou, Z.Q., Yang, Z.M., Gou, Z.B., DuFrane, S.A., 2014b. 1695 Increasing magmatic oxidation state from Paleocene to Miocene in the eastern Tibetan 1696 Gangdese belt: implication for collision-related porphyry Cu–Mo±Au mineralization. 1697 Econ. Geol. 109, 1943–1965.1698 Wang, R., Richards, J.P., Hou, Z.Q., Yang, Z.M., 2014c, Extent of underthrusting of the 1699 Indian plate beneath Tibet controlled of Miocene porphyry Cu-Mo±Au deposits. Miner. 1700 Deposita 49, 165–1731701 Wang, R., Richards, J.P., Zhou, L.M., Hou, Z.Q., Stern, R.A., Creaser, R.A., Zhu, J.J., 2015a, 1702 the Role of Indian and Tibetan lithosphere in spatial distribution of Cenozoic magmatism 1703 and porphyry Cu-Mo±Au deposits in the Gangdese belt, southern Tibet. Earth-Sci. Rev. 1704 150, 68–94.1705 Wang, R., Richards, J.P., Hou, Z.Q., An, F., Creaser, R.A., 2015b. Zircon U-Pb age and Sr-1706 Nd-Hf-O isotope geochemistry of the Paleocene–Eocene igneous rocks in western 1707 Gangdese: Evidence for the timing of Neo-Tethyan slab breakoff. Lithos, 224–225, 179–1708 194.1709 Wang, R., Collins, W.J., Weinberg, R.F., Li, J.X., Li, Q.Y., He, W.Y., Richards, J.P., Hou, 1710 Z.Q., Zhou, L.M., Stern, R.A., 2016. Xenoliths in ultrapotassic volcanic rocks in the Lhasa 1711 block: direct evidence for crust-mantle mixing and metamorphism in the deep crust. 1712 Contrib. Mineral. Petrol., 171, 621713 Wang, R., Collins, W.J, Weinberg, R.F, Richards, J.P., He, W.Y., 2017a, Reply to the 1714 comments on “Xenoliths in ultrapotassic volcanic rocks in the Lhasa block: direct 1715 evidence for crust–mantle mixing and metamorphism in the deep crust”. Contrib. Mineral. 1716 Petrol., 170:20. 1717 Wang, R., Tafti, R., Hou, Z.Q., Shen, Z.C., Guo, N., Evans, N., Jeon, H., Li, Q.Y., Li, W.K., 1718 2017b. Across-arc geochemical variation in the Jurassic magmatic zone, Southern Tibet: 1719 Implication for continental arc-related porphyry Cu-Au mineralization. Chem. Geol., 451, 1720 116–134.1721 Warren, C.J., Beaumont, C., Jamieson, R.A., 2008. Deep subduction and rapid exhumation: 1722 the role of crustal strength and strain weakening in continental subduction and ultrahigh 1723 pressure rock exhumation. Tectonics 27, TC6002. doi:10.1029/2008TC0022921724 Waters, L.E., Lange, R.A., 2015. An updated calibration of the plagioclase-liquid 1725 hygrometer-thermometer applicable to basalts through rhyolites. Am. Mineral. 100, 2172–1726 2184.1727 Watson, E.B., Harrison, T.M., 2005, Zircon thermometer reveals minimum melting

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1728 conditions on earliest Earth. Science 308, 841–844.1729 Weinberg, R.F. and Dunlap, W.J., 2000, Growth and deformation of the Ladakh Batholith, 1730 NW Himalayas: implications for timing of continental collision and origin of calc-alkaline 1731 batholiths. J. Geol. 108, 303–3201732 Weinberg, R.F., Hasalová, P., 2015, Water-fluxed melting of the continental crust: A review. 1733 Lithos 212–215, 158–188.1734 Wen, D.R., 2007. The Gangdese batholith, Southern Tibet. Ages, geochemical characteristics 1735 and petrogenesis. Unpublished Ph.D. thesis, National Taiwan University, 120 p.1736 Wen, D.R., Liu, D.Y., Chung, S.L., Chu, M.F., Ji, J.Q., Zhang, Q., Song, B., Lee, T.Y., Yeh, 1737 M.W., Lo, C.H., 2008. Zircon SHRIMP U-Pb ages of the Gangdese Batholith and 1738 implications for Neotethyan subduction in southern Tibet. Chem. Geol., 252, 191–201.1739 Wendlandt, R.F., Eggler, D.H., 1980. The origins of potassic magmas; 2, Stability of 1740 phlogopite in natural spinel lherzolite and in the system KAlSiO4-MgO-SiO2-H2O-CO2 at 1741 high pressures and high temperatures. Am. J. Sci. 280, 421–458.1742 Willet, S.D., Beaumont, C., 1994. Subduction of Asian lithospheric mantle beneath Tibet 1743 inferred from modesl of continental collision. Nature 369, 642–645.1744 Williams, H.M., 2000. Magmatic and tectonic evolution of Southern Tibet and the Himalaya: 1745 Ph.D. thesis, The Open University, 329 p.1746 Williams, H.M., Turner, S.P., Kelley, S.P., Harris, N.B.W., 2001. Age and composition of 1747 dikes in Southern Tibet: new constraints on the timing of east–west extension and its 1748 relationship to post-collisional volcanism. Geology 29, 339–342.1749 Williams, H.M., Turner, S.P., Pearce, J.A., Kelley, S.P., Harris, N.B.W., 2004. Nature of the 1750 source regions for post-collisional, potassic magmatism in southern and northern Tibet 1751 from geochemical variations and inverse trace element modelling. J. Petrol. 45, 555–607.1752 Williamson, B.J., Herrington, R.J., Morris, A., 2016, Porphyry copper enrichment linked to 1753 excess aluminium in plagioclase. Nature Geosci. 9, 237–241.1754 Wood, B.J., Turner, S.P., 2009. Origin of primitive high-Mg andesite: constraints from 1755 natural examples and experiments. Earth Planet. Sci. Lett. 283, 59–66.1756 Xu, B., Griffin, W.L., Xiong, Q., Hou, Z.Q., O’Reilly, S.Y., Guo, Z., Pearson, N.J., Gréau, Y., 1757 Yang, Z.M., Zheng, Y.C., 2016. Ultrapotassic rocks and xenoliths from South Tibet: 1758 Contrasting styles of interaction between lithospheric mantle and asthenosphere during 1759 continental collision. Geology doi: 10.1130/g38466.11760 Yang, Z.M., Hou, Z.Q., White, N.C., Chang, Z.S., Li, Z.Q., Song, Y.C., 2009. Geology of the 1761 post-collisional porphyry copper-molybdenum deposit at Qulong, Tibet: Ore Geol. Rev. 1762 36, 133–159.1763 Yang, Z.M., Lu, Y.J., Hou, Z.Q., Chang, Z.S., 2015. High-Mg diorite from Qulong in 1764 Southern Tibet: Implications for the genesis of Adakite-like intrusions and associated 1765 porphyry Cu deposits in collisional orogens. J. Petrol. 56, 227–254.1766 Yang, Z.M., Goldfarb, R., Chang, Z.S., 2016. Generation of postcollisional porphyry copper 1767 deposits in southern Tibet triggered by subduction of the Indian continental plate. Econ. 1768 Geol. S19, 279–300.1769 Ye, Z., Gao, R., Li, Q., Zhang, H., Shen, X., Liu, X., Gong, C., 2015. Seismic evidence for 1770 the North China plate underthrusting beneath northeastern Tibet and its implications for 1771 plateau growth. Earth Planet. Sci. Lett. 426, 109–117.1772 Yin, A., 2010. Cenozoic tectonic evolution of Asia: a preliminary synthesis. Tectonophysics, 1773 488, 293–325.1774 Yin, A., Harrison, T. M., 2000. Geologic evolution of the Himalayan-Tibetan orogen. Annu. 1775 Rev. Earth Planet. Sci. 28, 211–280.1776 Ying, L. J., Wang, C. S., Tang, J. X., Wang, D. H., Qu, W. J., Li, C., 2014. Re-Os 1777 systematics of sulfides (chalcopyrite, bornite, pyrite and pyrrhotite) from the Jima Cu–Mo

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1778 deposit of Tibet, China. J. Asian Earth Sci. 79, 497–506.1779 Zhang, Z.M., Zhao, G.C., Santosh, M., Wang, J.L., Dong, X., Shen, K., 2010. Late 1780 Cretaceous charnockite with adakitic affinities from the Gangdese batholith, southeastern 1781 Tibet: evidence for Neo-Tethyan mid-ocean ridge subduction? Gondwana Res., 17, 615–1782 631.1783 Zhao, J., Yuan, X., Liu, H., Kumar, Pei, S., Kind, R., Zhang, Z., Teng, J, Ding, L., Gao, X., 1784 Xu, Q., Wang, W., 2010. The boundary between the Indian and Asian tectonic plates 1785 below Tibet. PANS 107, 11229–11233.1786 Zhao, J. X., Qin, K. Z., Li, G. M., Li, J. X., Xiao, B., Chen, L., Yang, Y. H., Li, C., Liu, Y. S., 1787 2014. Collision-related genesis of the Sharang porphyry molybdenum deposit, Tibet: 1788 evidence from zircon U-Pb ages, Re-Os ages and Lu-Hf isotopes. Ore Geol. Rev. 56, 312–1789 326.1790 Zhao, W., Kumar, P., Mechie, J., Kind, R., Meissner, R., Wu, Z., Shi, D., Su, H., Xue, G., 1791 Karplus, M., Tilmann, F., 2011. Tibetan plate overriding the Asian plate in central and 1792 northern Tibet. Nat. Geosci.4, 870–873.1793 Zhao, Z.D., Mo, X.X., Dilek, Y, Niu, Y.L., DePaolo, D.J., Robinson P., Zhu, D.C., Sun, 1794 C.G., Dong, G.C., Zhou, S., Luo, Z.H., Hou, Z.Q., 2009. Geochemical and Sr-Nd-Pb-O 1795 isotopic compositions of the post-collisional ultrapotassic magmatism in SW Tibet: 1796 petrogenesis and implications for India-continental subduction beneath southern Tibet. 1797 Lithos 113, 190–212.1798 Zheng, W.B., Tang, J.X., Zhong, K.H., Ying, L.J., Leng, Q.F., Ding, S, Lin, B., 2016, 1799 Geology of the Jiama porphyry copper-polymetallic systems, Lhasa Region, China. Ore 1800 Geol. Rev. 74, 151–169.1801 Zheng, Y. C., Hou, Z. Q., Li, W., Liang, W., Huang, K. X., Li, Q. Y., Sun, Q. Z., Fu, Q., 1802 Zhang, S., 2012. Petrogenesis and geological implications of the Oligocene Chongmuda-1803 Mingze adakite-like intrusions and their mafic enclaves, southern Tibet. J. Geol. 120, 1804 647–669.1805 Zheng, Y.Y., Zhang, G.Y., Xu, R.K., Gao, S.B., Pang, Y.C., Cao, L., Du, A.D., Shi, Y.R., 1806 2007. Geochronologic constraints on magmatic intrusions and mineralization of the 1807 Zhunuo porphyry copper deposit in Gangdese, Tibet. Chin. Sci. Bull. 52, 3139–3147.1808 Zheng, Y. Y., Sun, X., Gao, S. B., Zhao, Z. D., Zhang, G. Y., Wu, S., You, Z. M., Li, J. D., 1809 2014. Multiple mineralization events at the Jiru porphyry copper deposit, southern Tibet: 1810 implications for Eocene and Miocene magma sources and resource potential. J. Asian 1811 Earth Sci. 79, 842–857.1812 Zheng, Y.C., Hou, Z.Q., Gong, Y.L., Liang, W., Sun Q.Z., Zhang, S., Fu, Q., Huang, K.X., 1813 Li, Q.Y., Li, W., 2014. Petrogenesis of Cretaceous adakite-like intrusions of the 1814 Gangdese plutonic belt, southern Tibet: Implications for mid-ocean ridge subduction and 1815 crustal growth. Lithos 190–191, 240–263. 1816 Zhou, S., Mo, X.X., Zhao, Z.D., Qiu, R.Z., Niu, Y.L., Guo, T.Y., Zhang, S.Q., 2010. 1817 40Ar/39Ar geochronology of post-collisional volcanism in the middle Gangdese Belt, 1818 southern Tibet. J. Asian Earth Sci., 37, 246–258.1819 Zhu, D.C., Mo, X.X., Zhao, Z.D., Niu, Y.L., Wang, L.Q., Chu, Q.H., Pan, G.T., Xu, J.F., 1820 Zhou, C.Y., 2010. Presence of Permian extension- and arc-type magmatism in southern 1821 Tibet: petrogeographic implications. Geol. Soc. Am. Bull. 122, 979–993.1822 Zhu, D.C., Zhao, Z.D., Niu, Y.L., Dilek, Y., Mo, X.X., 2011. Lhasa terrane in southern Tibet 1823 came from Australia. Geology 39, 727–730. 1824 Zhu, D.C., Zhao, Z.D., Niu, Y., Dilek, Y., Hou, Z.Q., Mo, X.X., 2013. Origin and pre-1825 Cenozoic evolution of the Tibetan plateau. Gondwana Res. 23, 1429–1454.1826 Zhu, D.C., Wang, Q., Zhao, Z.D., Chung, S.L., Cawood, P.A., Niu, Y.L., Liu, S.A., Wu, 1827 F.U., Mo X.X., 2015. Magmatic record of India-Asia collision. Sci. Rep. 5: 14289.

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1828 Zhu, D.C., Li, S.M., Cawood, P.A., Wang, Q., Zhao, Z.D., Liu, S.A., Wang, L.Q., 2016. 1829 Assembly of the Lhasa and Qiangtang terranes in central Tibet by divergent double 1830 subduction. Lithos 245, 7–171831 Zhu, D.C., Wang, Q., Cawood, P.A., Zhao, Z.D., Mo, X.X., 2017. Raising the Gangdese 1832 Mountains in southern Tibet. J Geophys. Res. Solid Earth 122, 214–223.18331834183518361837 Appendix-Analytical methods

1838 Zircon trace element analysis

1839 Here, we add new zircon chemistry results to compiled data in order to determine the

1840 magmatic temperatures, water contents, and oxidation states of the magmatic rocks. Zircons

1841 were mounted in epoxy with chips of standard ATHO-G and NIST 612 glass. A frequency

1842 quintupled Nd-YAG laser (UP 213, New Wave Research) was used to ablate the zircons and

1843 standard, and ablated material was carried by He-Ar gas (flow rates of 0.73 L He/min and

1844 0.85 L Ar/min) to a Finnigan MAET ELEMENT II high-resolution inductively coupled

1845 plasma mass spectrometer (HR-ICP-MS) at the National Research Center for Geoanalysis,

1846 Beijing, China. Laser ablation pits were approximately 40 μm wide and 20 μm deep. Data

1847 were acquired for 12 s with the laser off, and 43 s with the laser on, then the system was

1848 flushed with He-Ar gas for 15 s with the laser off. Each block of ten zircon analyses was

1849 bracketed by analysis of standard glass NIST 612, which was used to correct for mass bias

1850 drift during analysis (Pearce et al., 1997). The calibration procedure using internal standards

1851 and matrix normalization followed Hu et al. (2008). Calibration was conducted by

1852 normalizing count rates for each analysed element with Si to obtain its concentration, and

1853 assuming SiO2 to be stoichiometric in zircon (ZrSiO4) with a concentration of ca. 32.8 wt.%.

1854 Accuracy for selected elements, as determined by reproducibility of laboratory standards and

1855 duplicates, is within 10 relative %.

1856 In order to avoid mineral and glass inclusions, inherited cores, and fractures, the laser

1857 ablation spots in zircon were selected carefully. Apatite inclusions are common in zircon

37183719372037213722372337243725372637273728372937303731373237333734373537363737373837393740374137423743374437453746374737483749375037513752375337543755375637573758375937603761376237633764376537663767376837693770377137723773377437753776

65

1858 grains and can strongly affect REE contents, so phosphorus and calcium were included in the

1859 list of elements measured in order to monitor for apatite inclusions.

1860

1861 Electron microprobe analyses, plagioclase data

1862 Plagioclase in igneous rocks from the Gangdese belt were analysed by electron

1863 microprobe at the University of Alberta, Canada. Electron microprobe data were acquired on

1864 a JEOL 8900 instrument operated at 15 kV accelerating voltage and 15 nA probe current, and

1865 with a beam diameter of 1 to 5 μm. A variety of minerals, oxides, and elemental standards

1866 were used for calibration, and data reductions were undertaken with the CITZAF routine of

1867 J.T. Armstrong (as implemented by P. Carpenter in the JEOL software). The limits of

1868 detection are typically lower than 500 ppm (≤800 ppm for Mn, Fe, Ti, Cr, and Ni; ≤5000 ppm

1869 for F), and analytical precision for major elements is better than one relative percent.

1870 The Al composition of plagioclase can be used to estimate magmatic water contents

1871 (Lange et al., 2009; Waters and Lange, 2015; Williamson et al., 2016). We analysed EPMA

1872 compositions of plagioclase from Paleocene-Eocene Gangdese arc intrusions, Miocene PCD-

1873 bearing intrusions, and Miocene volcanic rocks, and combined these data with analyses of

1874 plagioclase from the Qulong deposit (Xiao, 2011). These data are listed in Table A3.

1875

1876 Figure A1 Two profiles of the geometry of the colliding lithospheres discussed in Zhao et al.

1877 (2010): (A) Underthrusting of Indian lithosphere along the western profile along longitude

1878 ~80°E; (B) northward subduction of Indian lithosphere along the eastern profile along

1879 longitude ~87°E. Abbreviations: BNS = Bangong-Nujiang suture; MBT = Main central thrust;

1880 YZS = Indus-Yarlung Tsangpo suture.

1881

37773778377937803781378237833784378537863787378837893790379137923793379437953796379737983799380038013802380338043805380638073808380938103811381238133814381538163817381838193820382138223823382438253826382738283829383038313832383338343835

66

1882 Figure A2 Trace element composition of zircons from the six main Cenozoic suites in the

1883 Gangdese belt: (A) Ce4+/Ce3+ vs. Eu/Eu*; (B) Th/U vs. T(ti-zr); (C) Ce4+/Ce3+ vs. T(ti-zr); (D)

1884 Ce4+/Ce3+ vs. Th/U. References for data provided in Table A2.

1885

1886 Figure A3 Chondrite normalized REE diagram for zircons from the six main Cenozoic suites

1887 in the Gangdese belt: (A) W Miocene trachytes, E Miocene trachytes, and E Miocene

1888 transitional monzonites; (B) Paleocene-Eocene Gangdese arc calc-alkaline rocks, W Miocene

1889 high-Sr/Y granitoids, and E Miocene high-Sr/Y granitoids. References for data provided in

1890 Table A1.

1891

1892 Figure A4 Trace element composition of zircons from Paleocene-Eocene Gangdese arc

1893 igneous rocks: (A) T(ti-zr) vs. age; (B) Eu/Eu*vs. age; (C) Ce4+/Ce3+ vs. age. Note that

1894 zircons from 53–50 Ma igneous rocks show significantly higher T(ti-zr) =800–900 °C

1895 compared to older igneous rocks (<800C), but lower Eu/Eu* and Ce4+/Ce3+ ratios indicating

1896 input of mantle-derived melts with low oxidation states. References for data provided in

1897 Tables A1, A2.

1898

1899 Table A1. Published major and trace element compositions of Cenozoic igneous rocks from

1900 the Gangdese Belt

1901

1902 Table A2. Zircon trace element compositions for Cenozoic igneous rocks from the Gangdese

1903 Belt..

1904

1905 Table A3. EPMA compositions for plagioclase from Cenozoic igneous rocks from the

1906 Gangdese Belt.

38363837383838393840384138423843384438453846384738483849385038513852385338543855385638573858385938603861386238633864386538663867386838693870387138723873387438753876387738783879388038813882388338843885388638873888388938903891389238933894

80°

86°

84°

82°

88° 90° 94°92°

31°

30°

29°

32°

Shiquan river-Nam Tso Mélange

Luobadui-Milashan fault.Northern Lhasa subterrane

Southern Lhasa subterrane

Central Lhasa subterrane

Gar

Hor

Xungba

Yanhu

Gerze

Nyima

Saga Sangsang

Xigaze

Luoza

Bayi

Nang0 50 100 km

AwengTso

Damxung

Xainza

Geji

Lhasa

ASIA

INDIA

Miocene trachyte suite

Cretaceous batholithsLinzizong volcanic successions

Paleocene–Eocene intrusions

Oligocene granitoids (sparse)Miocene high-Sr/Y granitoidsMiocene transitional monzonites (sparse)

Proterozoic strata

Cambrian strata

City

Suture or fault

Graben basin

Namling

Yangying

Gongbogyamda

Mamba

NDazi

Bangong-Nujiang suture

Porphyry deposit

Suojin

Qiangtang terrane

Zedang

Indus-Yarlung Tsangpo suture

Qulong

Bangpu Sharang

Jiama

Chongjiang

Tangbula

Tinggong

Jiru

Xietongmen

Zhunuo

Mantle upwelling

Indian crust

high-Sr/Y melts

km

60

0

120

Indian crust

0 25 50 km Asthenosphere

Asthenosphere

Thickened Tibetan crust

Thickened Tibetan crust

Arc cumulates

Indian lithosphere

Indian lithosphere

TibetanSCLM

Continental underthrusting (flat subduction) (Western Gangdese)

Continental subduction (Eastern Gangdese): Indian SCLM is source of trachytic melts

Arc cumulates

Tibetan SCLM

IYS

IYSkm

60

0

120

750°C

750°C

1100°C

1100°C

1100°C

750°C750°C

1100°C

geotherm

A

B

trachytes

TrachytesTrachytes

Trachytes

horizontal scale

0 25 50 kmhorizontal scale

Himalayan orogen

Himalayan orogen

Asthenosphere

leucogranite

Linzizong suite

Gangdese batholith

Linzizong suiteGangdese batholith

high-Sr/Y granitoids

oxidized carbonate and evaporitic sediments

trachytic melts

Mantle upwelling

Indian crust

Asthenosphere

Thickened Tibetan crust

Indian lithosphere

Continental subduction (Eastern Gangdese): underthrusted Tibetan SCLM is source of trachytic melts

Arc cumulates

Tibetan SCLM

IYS

km

60

0

120

750°C750°C

1100°C

1100°C

CTrachytes

0 25 50 kmhorizontal scale

Himalayan orogen

Asthenosphere

Linzizong suite

Gangdese batholith

fore-arc SCLM