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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).
650651652653654655656657658659660661662663664665666667668669670671672673674675676677678679680681682683684685686687688689690691692693694695696697698699700701702703704705706707708
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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
768769770771772773774775776777778779780781782783784785786787788789790791792793794795796797798799800801802803804805806807808809810811812813814815816817818819820821822823824825826
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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
827828829830831832833834835836837838839840841842843844845846847848849850851852853854855856857858859860861862863864865866867868869870871872873874875876877878879880881882883884885
16
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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24
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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28
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
18301831183218331834183518361837183818391840184118421843184418451846184718481849185018511852185318541855185618571858185918601861186218631864186518661867186818691870187118721873187418751876187718781879188018811882188318841885188618871888
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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35
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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36
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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37
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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39
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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41
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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42
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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43
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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44
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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45
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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46
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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47
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
30103011301230133014301530163017301830193020302130223023302430253026302730283029303030313032303330343035303630373038303930403041304230433044304530463047304830493050305130523053305430553056305730583059306030613062306330643065306630673068
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