Permian ultrahigh–temperature reworking in the southern Chinese Altai: Evidence from petrology, P–T estimates, zircon and monazite U–Th–Pb geochronology

Journal Pre-proof Permian ultrahigh–temperature reworking in the southern Chinese Altai: Evidence from petrology, P–T estimates, zircon and monazite U–Th–Pb geochronology Zhao Liu, Omar Bartoli, Laixi Tong, Yi–Gang Xu, Xiaolong Huang PII:

S1342-937X(19)30248-5

DOI:

https://doi.org/10.1016/j.gr.2019.08.007

Reference:

GR 2204

To appear in:

Gondwana Research

Received Date: 15 January 2019 Revised Date:

26 July 2019

Accepted Date: 6 August 2019

Please cite this article as: Liu, Z., Bartoli, O., Tong, L., Xu, Y.–G., Huang, X., Permian ultrahigh– temperature reworking in the southern Chinese Altai: Evidence from petrology, P–T estimates, zircon and monazite U–Th–Pb geochronology, Gondwana Research, https://doi.org/10.1016/j.gr.2019.08.007. This is a PDF file of an article that has undergone enhancements after acceptance, such as the addition of a cover page and metadata, and formatting for readability, but it is not yet the definitive version of record. This version will undergo additional copyediting, typesetting and review before it is published in its final form, but we are providing this version to give early visibility of the article. Please note that, during the production process, errors may be discovered which could affect the content, and all legal disclaimers that apply to the journal pertain. © 2019 International Association for Gondwana Research. Published by Elsevier B.V. All rights reserved.

1

Permian ultrahigh–temperature reworking in the southern Chinese Altai:

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Evidence from petrology, P–T estimates, zircon and monazite U–Th–Pb

3

geochronology

4

Zhao Liua, b, c, Omar Bartolic, Laixi Tongd, *, Yi–Gang Xua, Xiaolong Huanga

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a

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Academy of Sciences, Guangzhou 510640, China

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b

University of Chinese Academy of Sciences, Beijing 100049, China

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c

Dipartimento di Geoscienze, Università di Padova, Via Gradenigo 6, 35131 Padua, Italy

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d

State Key Laboratory of Continental Dynamics, Department of Geology, Northwest University,

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Xi’an 710069, China

State Key Laboratory of Isotope Geochemistry, Guangzhou Institute of Geochemistry, Chinese

13 14 15 16 17 18

*

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E–mail address: [email protected] (L. Tong).

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Corresponding author. Tel: +86 29 88302312; Fax: +86 29 88302202.

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ABSTRACT

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The Chinese Altai orogen formed in the Paleozoic is an important part of the Central Asian

26

Orogenic Belt (CAOB), and the study on the metamorphism will provide novel and robust

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constraints on its tectonic evolution. In this study, we investigate our newly recognized garnet–

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orthopyroxene–cordierite granulites at Wuqiagou area in the southern Chinese Altai. Detailed

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petrographic study and P–T estimates suggest four distinct metamorphic stages of mineral

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assemblages: (1) pre–peak (M1) stage containing the spinel–cordierite–bearing association or

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biotite–plagioclase–quartz–bearing inclusion–phase assemblage, with P–T conditions of 3.0–4.0

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kbar/700–750 °C; (2) peak ultrahigh–temperature (UHT) (M2) stage represented by relatively

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coarse–grained garnet–orthopyroxene–cordierite–bearing porphyroblastic assemblage, with high–

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Al2O3 contents (up to ~8.7 wt.%) in orthopyroxene and P–T conditions of ∼8.0 kbar/~980 °C; (3)

35

post–peak high–temperature granulite facies (M3) stage consisted of orthopyroxene–cordierite and

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cordierite–quartz corona assemblages, formed during cooling and moderate decompression; and (4)

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post–peak upper amphibolite facies (M4) stage represented by retrograde biotite–plagioclase–quartz

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intergrowths. These four discrete metamorphic stages define an anticlockwise P–T path involving a

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post–peak moderate decompression followed by nearly isobaric cooling process. LA–ICP–MS U–

40

Pb age dating results of metamorphic zircons for UHT samples show two weighted mean ages of

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~390 Ma and ~280 Ma. We propose that the M1 stage might occur in the middle Devonian, whereas

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the near–peak UHT stage probably occurred in the early Permian. The Permian UHT

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metamorphism was further supported by the monazite U–Th–Pb dating results (287.9 ± 2.1 Ma),

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reflecting a prominent HT–UHT reworking event in the late Paleozoic. We proposed that the

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Permian–age UHT reworking event in the southern Chinese Altai probably occurred in a post–

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orogenic or intraplate extensional tectonic setting associated with the input of external heat, related

47

to the underplating of deep–derived magma as a result of the Tarim mantle plume activity.

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Keywords: UHT metamorphism; Chinese Altai; P–T path; Geochronology; Mantle plume

50 51

1. Introduction

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Ultrahigh temperature (UHT) metamorphism is the most thermally extreme type of crustal

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metamorphism, with temperatures exceeding 900 °C at moderate pressure (7–13 kbar; 20–40 km)

55

(Kelsey, 2008; Harley, 2008; Santosh et al., 2012). The study of UHT terranes can provide

56

important insights into the formation and evolution of deep continental crust (Kelsey and Hand,

57

2015; Korhonen et al., 2014). Some mineral assemblages in rocks of highly aluminous and

58

magnesian bulk composition (Mg–, Al–rich granulites), including assemblages such as sapphirine +

59

quartz, orthopyroxene (> 8.0 wt.% Al2O3) + sillimanite ± quartz, low Zn spinel + quartz and

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osumilite + garnet, are the key indicators of such extreme metamorphic conditions (Harley, 1998,

61

2008; Tsunogae et al., 2011; Santosh et al., 2012). But under highly oxidizing or some other certain

62

conditions, even these assemblages may be stabilized below 900 °C (Kelsey et al., 2008).

63

Conventional geothermometers, which are mainly based on Fe–Mg exchange reactions, generally

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yield erroneous low temperature estimates due to post–peak diffusional cation exchange (e.g.,

65

Harley, 1989). This drawback, however, could be potentially overcome by using alternative

66

approaches such

as

garnet–orthopyroxene thermobarometer based

on Al–solubility in

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orthopyroxene (Pattison et al., 2003), two–feldspar thermometry (Jiao et al., 2011), Ti–in–zircon

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thermometry (Ferry and Watson, 2007), and Zr–in–rutile thermometry (Tomkins et al., 2007).

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A long–term dispute exists on the mechanisms which make crust extremely hot with

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geothermal gradient of ≥ 20 °C km–1 (Brown, 2007; Kelsey, 2008; Santosh and Kusky, 2010), to

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produce UHT granulites. Although about 60 UHT granulite terranes of Neoarchean to Miocene

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have been identified so far (Brown, 2007; Kelsey, 2008; Pownall et al., 2014; Kelsey and Hand,

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2015), only few UHT metamorphic events are considered to have occurred in the past 500 Ma (e.g.,

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Nam et al., 2001; Zhao et al., 2010; Galli et al., 2011; Li et al., 2014; Pownall et al., 2014; Tong et

75

al., 2014; Zhang et al., 2015b). Underplating of hot mafic magmas (Guo et al., 2012; Li et al., 2014;

76

Tong et al., 2014) or exhumation of subcontinental lithospheric mantle (Pownall et al., 2014) have

77

been proposed as heat sources for UHT metamorphism of crustal rocks.

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The Chinese Altai orogenic belt formed in the Paleozoic is an important part of the Central

79

Asian Orogenic Belt (CAOB), accompanying with remarkable metamorphism (e.g., Li et al., 2004,

80

2014; Wei et al., 2007; Wang et al., 2009b; Tong et al., 2014; Broussolle et al., 2018). The early

81

accretionary wedge in the Chinese Altai pervasively experienced a middle Devonian tectono–

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metamorphic event (e.g., Wei et al., 2007; Jiang et al., 2010). The Devonian orogenic architecture

83

was subsequently reworked by the HT–UHT metamorphism on its southern margin (e.g., Wang et

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al., 2009b; Tong et al., 2013, 2014; Jiang et al., 2015; Broussolle et al., 2018). Contrasting P–T

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paths and tectonic models were proposed to explain the late Paleozoic UHT metamorphism in the

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southern Chinese Altai:

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(i) Altai UHT granulites followed a clockwise P–T path (Li et al., 2010; Yang et al., 2015b), with a

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first isothermal decompression (ITD) and a second isobaric cooling (IBC) retrograde P–T paths

89

(Yang et al., 2015b). Li et al. (2010) correlated the UHT metamorphism with the collisional

90

orogeny between Siberia and Kazakhstan–Junggar plate.

91

(ii) Altai UHT granulites from Wuqiagou area were exhumed along an anticlockwise P–T trajectory

92

with a post–peak near–ITD process at 5–6 kbar/890–940 °C (Li et al., 2014). Slab break–off

93

which caused asthenospheric upwelling and heat flux at 320–290 Ma might contribute to Altai

94

UHT metamorphism (Li et al., 2014; Yang et al., 2015b);

95

(iii) P–T conditions of the Altai UHT granulites from Kalasu area define an anticlockwise P–T path

96

of initial prograde heating and increase in pressure followed by a post–peak near–IBC process

97

(Tong et al., 2014). Underplating and heating of mantle–derived mafic magma as a result of the

98

Tarim mantle plume might provide the heat flux necessary for the Permian HT–UHT

99

metamorphism (Wang et al., 2014; Tong et al., 2013, 2014; Yang et al., 2015a; Liu and Tong,

100

2015).

101

Accordingly, the P–T path of the Altai UHT granulites is not well constrained and the

102

tectonothermal evolution of the southern Chinese Altai is still ambiguous and debated. In this study,

103

we investigate the petrology, mineral chemistry, zircon and monazite geochronology, and P–T

104

trajectory of garnet–orthopyroxene–cordierite–bearing UHT granulites at Wuqiagou area. UHT

105

conditions are retrieved from the high alumina contents of orthopyroxene (~8.7 wt.% Al2O3),

106

conventional geothermobarometry calculations and preliminary phase equilibria modelling. The P–

107

T path recorded in the UHT metapelitic granulites will place important constraints on the

108

continental crustal growth and tectonothermal evolution of the southern Chinese Altai in the

109

Paleozoic.

110

111

2. Geological background

112 113

The Altai orogen is an important part of the CAOB (Jahn et al., 2004). The NW–SE trending

114

Chinese Altai orogenic belt is bounded by the Siberian plate to the north and the Kazakhstan–

115

Junggar plate to the south (Windley et al., 2007). The Chinese Altai orogen comprises various

116

lithological types, mainly including volcanic, pyroclastic and metasedimentary rocks, high–grade

117

metamorphic rocks and large amounts of granitoids (Windley et al., 2002; Jiang et al., 2015).

118

Five fault–bounded terranes have been identified based on the stratigraphy, metamorphism,

119

deformation patterns and chronology (Fig. 1) (Windley et al., 2002; Wang et al., 2006, 2009a).

120

Terrane Ⅰ consists mainly of the late Devonian to the early Carboniferous meta–clastic rocks and

121

limestone intercalated with minor arc–like volcanic rocks. Terrane Ⅰ is composed mainly of the

122

Neoproterozoic to the middle Ordovician sedimentary and volcanic rocks of the Habahe Group

123

(Yuan et al., 2007), which experienced lower greenschist facies metamorphism. Terrane Ⅰ in the

124

central part of the Altai orogenic belt is the largest one and is composed mainly of the early Silurian

125

and the early Devonian flysch sequence of the Habahe Formation (Long et al., 2010), among which,

126

the minor ~502 Ma felsic volcanic rocks which experienced greenschist– to upper amphibolite

127

facies metamorphism have been interpreted to represent components of a continental arc (Windley

128

et al., 2002; Yang et al., 2011). Terrane Ⅰ consists of the late Silurian to the early Devonian arc–like

129

volcanic and pyroclastic rocks in the lower part and the middle Devonian turbidites and pillow–

130

basalts in the upper part, showing a spectrum of metamorphic zones from greenschist to upper

131

amphibolite and locally granulite facies conditions (Wang et al., 2009b; Tong et al., 2014). Terrane

132

Ⅰ is bounded by the Erqis fault fault in the south. It includes a complex sequence including a

133

possible Precambrian basement, the early Paleozoic–Devonian sediments and the late

134

Carboniferous volcanoclastics, metamorphosed under greenschist to amphibolite facies conditions.

135

Mafic granulites and UHT pelitic granulites were reported from this terrane (Li et al., 2004, 2010,

136

2014; Chen et al., 2006; Yang et al., 2015b; Liu and Tong, 2015). Rocks in the Junggar plate (south

137

of the Erqis fault belt) are dominated by the Devonian–Carboniferous volcanoclastics, which have

138

been metamorphosed to greenschist facies.

139

The tectonic evolution of the Chinese Altai orogen mainly involves five stages based on

140

previous studies (e.g., Windley et al., 2002, 2007; Wang et al., 2006; Niu et al., 2006; Jiang et al.,

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2010; Yang et al., 2011; Zhang et al., 2012; Wang et al., 2009b, 2014; Tong et al., 2014; He et al.,

142

2018): (i) a passive continental margin or peri–Gondwana terrane during the Neoproterozoic–early

143

Paleozoic; (ii) the development of a late Silurian to early Devonian arc environment related to the

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northward subduction of Junggar plate; (iii) continent–arc collision, subduction of ridge, or the

145

development of a possible back–arc basin in the middle to late Devonian; (iv) Permian (300–260

146

Ma) post–orogenic setting with a possible overprinting by the Tarim mantle plume; and (v)

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intraplate magmatism beginning in the Jurassic.

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High–grade gneissic rocks crop out extensively in the terranes Ⅰ and Ⅰ and are currently

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assigned to the Kemuqi and Fuyun Groups. Their presumed Precambrian age is widely accepted by

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many workers and led to the proposal of an Altai–Mongolia Precambrian basement or

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microcontinent (Windley et al., 2002; Li et al., 2006). However, the above interpretation is not

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supported by SHRIMP zircon U–Pb data for high–grade metamorphic rocks from the southeastern

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part of the Chinese Altai, which, instead, suggested that these high–grade rocks metamorphosed in

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the Paleozoic (Chen et al., 2006; Wang et al., 2009b; Jiang et al., 2010; Li et al., 2014; Tong et al.,

155

2014).

156

Two metamorphic events have been documented in the Chinese Altai orogenic belt (Broussolle

157

et al., 2018). The first tectono–metamorphic cycle was dated at 390–365 Ma (e.g., Zhuang, 1994;

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Hu et al., 2002; Zheng et al., 2007; Jiang et al., 2010), and considered to be linked with two distinct

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metamorphic stages characterized by M1 Barrovian–type MT–MP and M2 Buchan–type HT–LP

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field gradients, with metamorphic degrees ranging from greenschist– to amphibolite facies

161

conditions (up to 750 °C; Jiang et al., 2010, 2015). In terms of the tectonic process in the Devonian,

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researchers have envisaged models of active continental margin (He et al., 2018), arc–continent

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collision (Windley et al., 2002; Wei et al., 2007), back–arc spreading (Wang et al., 2006), slab

164

break–off (Niu et al., 2006) or ridge–subduction and the development of slab–window (Windley et

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al., 2007; Jiang et al., 2010).

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Recently, many HT–UHT metamorphic rocks were reported from the southern Chinese Altai,

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which mainly consisted of HT–UHT metapelitic granulites (Wang et al., 2009b; Li et al., 2014;

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Tong et al., 2014; Liu and Tong, 2015; Yang et al., 2015b), mafic granulites (Li et al., 2004; Chen

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et al., 2006; Liu and Tong et al., 2015), calc–silicate granulites (Yang et al., 2015a). The Permian

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high–grade rocks are mainly cropping out along NW–SE trending zone in the terranes Ⅰ and Ⅰ

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(e.g., Tong et al., 2014; Liu and Tong, 2015; Broussolle et al., 2018). Zircon U–Pb and monazite

172

U–Th–Pb geochronology of Altai HT–UHT granulites and gneisses yielded metamorphic ages of

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293–260 Ma (Chen et al., 2006; Briggs et al., 2007; Zheng et al., 2007; Wang et al., 2009b; Li et al.,

174

2014; Tong et al., 2014), which were interpreted by Broussolle et al. (2018) as the second tectono–

175

metamorphic cycle in the Chinese Altai orogeny.

176

Our samples were collected at Wuqiagou area (Fig. 2). The UHT granulites occur as lenses

177

within garnet–biotite–plagioclase gneiss, with the country rocks generally intruded by granitic dikes.

178

The Altai UHT granulites experienced partial melting as testified by migmatitic appearance (Fig. 3).

179 180

3. General petrography and reaction history

181 182

The Altai UHT granulite samples typically preserve a porphyroblastic texture composed of

183

garnet (12–15%), orthopyroxene (15–20%), cordierite (13–15%), sillimanite (~3%), spinel (~5%),

184

biotite (15–20%), plagioclase (10–12%), quartz (20–25%), with accessory anorthite, zircon,

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monazite, apatite, rutile and Fe–Ti oxides (< 5%).

186 187

M1 stage: interpretation of the prograde evolution

188

Garnet, orthopyroxene and cordierite porphyroblasts preserve a variety of single–phase or

189

multiphase mineral inclusions, which are ascribed to the pre–peak assemblages. These inclusions

190

show various sizes and shapes, generally with no significant preferred orientation. Tiny composite

191

grains of spinel + cordierite ± sillimanite are present in garnet and cordierite porphyroblasts (Figs.

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4b and c). Spinel, in addition to occurring as inclusions preserved in garnet, is also found in the

193

matrix, and generally associated with cordierite, fibrous sillimanite, biotite, Fe–Ti oxides and/or

194

anorthite (Figs. 4a and b). The spl1 + crd1 assemblage might reach equilibrium by the following

195

reaction (Bindu, 1997; White et al., 2007):

196

Bi + Sill ± Qtz → Sp1 + Crd1 ± Ksp + Liq (1)

197

The spinel–bearing assemblages were most likely to be formed in the pre–peak stage because

198

they are occasionally preserved as inclusions in garnet. Spinel in direct contact with quartz was

199

never observed, which is different from the nearby UHT granulites reported by Li et al. (2014).

200

Further inclusions in garnet porphyroblasts are biotite, plagioclase, quartz, Fe–Ti oxides and more

201

rarely, intergrowths of cordierite + magnetite ± biotite (Fig. 4b). Porphyroblastic orthopyroxene

202

generally encloses biotite, plagioclase, quartz and Fe–Ti oxides (Figs. 4e and f).

203 204

In conclusion, pre–peak (M1) assemblage mainly comprised garnet cores, spinel (spl1), cordierite (crd1), biotite (bi1), plagioclase (pl1), sillimanite, quartz and Fe–Ti oxides.

205 206

M2 stage: peak metamorphic assemblage

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Garnet (grt2), orthopyroxene (opx2) and texturally equilibrated cordierite (crd2) porphyroblasts

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probably represent the peak metamorphic assemblage. Magnetite, rutile and minor ilmenite are

209

present in the rock matrix (Figs. 4d, 5e and f). As described above, the growth of garnet and

210

orthopyroxene might document progress of the following reaction (Vielzeuf and Montel, 1994):

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Bi1 + Pl1 + Qtz1 → Grt2 + Opx2 ± Ksp + Liq (2)

212

Cuspate to lobate grains are present in leucocratic portions (Fig. 4g). These microstructures are

213

considered to be melt pseudomorphs (Holness and Sawyer, 2008; Sawyer, 2008) and are indicative

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of crystallization of localized melts among mineral boundaries. Some polycrystalline inclusions

215

occur isolated or are localized in clusters within garnet porphyroblasts (Fig. 4h). They often display

216

a round to negative crystal shape, range from 1 to 8 µm in diameter and contain multiple daughter

217

crystals of biobite, plagioclase, quartz and ilmenite (Figs. 4i and j), resembling the melt inclusions

218

(MI) described in peritectic garnet of other migmatitic and granulitic terranes (e.g., Acosta-Vigil et

219

al., 2010; Cesare et al., 2015; Bartoli et al., 2013, 2014, 2016). The observed inclusions would

220

represent an additional evidence of the former presence of melts in these rocks.

221

The garnet porphyroblasts are typically pale pink in colour and 0.7–2.0 mm across (Figs. 4b, 5c

222

and g). Orthopyroxenes are generally pale grayish to dark brown, and occur as anhedral grains and

223

have sizes ranging from 0.1 to 2.5 mm (Figs. 4e, f, 5a, e and f).

224

Cordierite grains vary in size from 0.02 to 2 mm and locally contain spinel, tabular biotite and

225

sillimanite needles (Figs. 4a and c). The peak cordierite might represent the only hydrous phase

226

during the peak metamorphic stage, suggesting a notably dry composition.

227 228

M3 stage: corona textures formation

229

Opx2 is extensively replaced by corona textures, which are mainly composed of orthopyroxene

230

(opx3), fine–grained cordierite (crd3) and Fe–Ti oxides (Figs. 5a and b). These opx3–crd3

231

symplectitic rinds are interpreted to be formed subsequently to the peak metamorphic stage and they

232

probably are the result of the release of the Mg–Tschermak’s component from opx2 according the

233

following retrograde reaction (Brandt et al., 2003):

234

High–Al Opx2 + Qtz → low–Al Opx3 + Crd3 (3; Fig. 6a)

235

Garnet may be mantled by cordierite moats (crd3), tiny biotite flakes and vermiform quartz,

236

with a distinct ‘spongy’ appearance (Figs. 5c and d). These microstructures likely reflect the melt–

237

consuming reaction (Cenki et al., 2002):

238

Grt + Liq → Crd + Bt + Qtz (4)

239

Formation of the retrograde opx3–crd3 corona texture was interpreted to be resulted from

240

decompression (Brandt et al., 2003). The melt consuming reaction 4 requires a combination of

241

cooling and decompression, which is consistent with an uplift path (Cenki et al., 2002).

242

243

M4 stage: formations of retrograde garnet, biotite, plagioclase and quartz

244

Retrograde platy and needle–like biotite (bi4), fine–grained plagioclase (pl4) and quartz might

245

form subsequently to the M3 stage around garnet and orthopyroxene porphyroblast rims (Figs. 5f

246

and g). The succession of M3 and M4 reactions (i.e., the M4 stage is likely to have occurred after

247

M3) is clearly visible in Figure 5f, where bi4–pl4–qtz intergrowths are locally grown around opx3 +

248

crd3 symplectites. Bi4–pl4–qtz intergrowths were generally considered to be resulted from

249

interaction of the melt with minerals, suggesting the progress of the following retrograde bi–

250

forming hydration reaction (Brandt et al., 2003; Holness et al., 2011):

251

Opx (Grt) + Liq → Bi4 + Pl4 + Qtz. (5)

252

Symplectitic opx3 is locally overgrown by fine–grained anhedral garnet (< 0.1 mm) (Fig. 5e).

253

These small garnets locally contain inclusions of opx3, indicating the growth of a second generation

254

of garnet at the expense of orthopyroxene. This garnet growth is usually considered to be related to

255

cooling (Brandt et al., 2003) linked to the following reaction (Harley, 1989):

256

Opx + Crd → Grt4 + Qtz. (6; Fig. 6a)

257

In places, some rutile needles were also observed in garnet (Fig. 5h). These rutile needles were

258

probably precipitated during M4 stage or further cooling under subsolidus conditions.

259 260

4. Whole–rock geochemistry and mineral chemistry

261 262

4.1. Whole–rock geochemistry

263 264

Major element oxides (wt.%) were determined on fused glass disks with a 1:8 sample to

265

Li2B4O7 flux ratio, using a Rigaku ZSX100e X–ray fluorescence (XRF) spectrometer in the Key

266

Laboratory of Isotope Geochronology and Geochemistry, Guangzhou Institute of Geochemistry.

267

The accuracy of the analyses is within 1 % for most major elements. Sample preparation techniques

268

and other details of procedures are described in the reference (Long et al., 2011). The geochemical

269

data are presented in Table 1.

270

The Altai UHT metapelitic granulites show variable SiO2 contents (51.42–56.49 wt.%). They

271

typically have moderate Al2O3 and MgO contents, with A/AFM (Al/(Al + Fetotal + Mg)) and Mg#

272

values of 0.32–0.36 and 52.03–52.77, respectively. The K2O + Na2O contents are in the range of

273

2.35–2.56 wt.%, with K2O higher than Na2O. P2O5 and MnO contents are negligible (< 0.39 wt.%).

274

The LOI (= loss on ignition) is normally ranging from 1.10 to 1.32 wt.%. In S(SiO2)–A(Al2O3 +

275

Fe2O3)–FM(FeO + MgO) diagram, they are plotted in the grt–opx–crd–qtz field (Fig. 6), which is

276

consistent with the major minerals observed in these samples.

277 278

4.2. Mineral chemistry

279 280

Mineral compositions were analyzed with a JXA–8100 microprobe at State Key Laboratory of

281

Isotopic Geochemistry, Guangzhou Institute of Geochemistry, Chinese Academy of Sciences, with

282

an accelerating voltage of 15 kv, a beam current of 3x10–8 Å, a beam width of 1 µm, and data

283

correction by using a ZAF method. Structural foumulae are given for fixed oxygen values with Fe3+

284

calculated by stoichiometric charge balance. Representative mineral compositions for the studied

285

samples are listed in Table 2 and Table 3.

286

Garnet is essentially an almandine–pyrope solid solution with minor spessartine and grossular

287

components (Alm61–72Py31–17Sps5–8Grs3–3) (Fig. 6). Slight chemical variations are observed for

288

garnets of distinct generations. The garnet porphyroblasts exhibit an obvious compositional zoning,

289

with a typically “M–like” profile from the core to rim (Figs. 6d and f). In details, the garnets display

290

a relatively low pyrope content in the core (Alm68–64Py22–28Sps7–4Grs3–4) and a slight outward

291

increase in Mg# from core (Mg# = 24.39) to mantle (Mg# = 32.90–33.85). Thereafter, the

292

composition becomes less magnesian rimward to Alm70–71Py18–18Sp7–8Grs5–3 (Mg# = 20.22–20.45).

293

Furthermore, the low Mg# contents of garnet rims are probably associated with the production of the

294

neighboring symplectites of bi4–pl4–qtz. Late grt4 grains have a much stricter compositional

295

variation of Alm65–68Py25–22Sps7–7Grs3–3.

296

Orthopyroxene from different positions has remarkably different Al2O3 contents. The core of

297

the orthopyroxene porphyroblast (Opx2) occasionally preserves markedly higher Al2O3 contents (up

298

to ~8.7 wt.%) than those of its rim (5.48–7.25 wt.%), with the corresponding XAl ( = Al/2) values of

299

0.19 and 0.12–0.16 in the formula, respectively. Symplectitic orthopyroxene (Opx3) shows wider

300

and lower Al2O3 contents (1.94–5.50 wt.%). However, they have similar XMg values ranging from

301

0.57 to 0.64.

302

Spinel inclusions preserved in garnet have moderate ZnO contents (~2.33 wt.%), relatively

303

higher Cr2O3 contents (~0.42 wt.%), and the highest XMg values (~0.34). The matrix spinel grains

304

contain ZnO of 0.82–3.25 wt.%, Cr2O3 of 0.15–0.39 wt.% and XMg of 0.21–0.27. The calculated

305

Fe2O3 contents of spinel by stoichiometric charge balance in this study generally vary from 1.83 to

306

2.86 wt.%, but occasionally as low as 0.36 wt.%, indicative of lower oxygen fugacity.

307

In all analyzed samples, the total sum of oxides in the cordierites is much lower than 100%,

308

suggesting the presence of H2O + CO2 in the structure. They have highly magnesian compositions

309

with XMg values ranging from 0.76 to 0.85. Cordierite included in garnet has the highest XMg value

310

of ~0.85. Peak Crd2 has relatively lower XMg values (0.76–0.78) than Crd1 (~0.80) stable with

311

spinel. Crd3 in the symplectitic intergrowths of opx3–crd3 has a typical XMg value of ~0.81 and crd3

312

moats around garnets have XMg values of 0.80–0.84.

313

Biotite preserved in garnet porphyroblasts has higher XMg values (0.66–0.72) and lower TiO2

314

contents (1.5–3.2 wt.%) than biotite in the spinel–bearing domains (XMg = 0.55–0.58, TiO2 = 4.80–

315

5.00). The matrix biotite shows XMg values of 0.59–0.65 and TiO2 contents of 1.81–3.35. Biotite

316

flakes display an increase in XMg values from the core (~0.61) to the rim (~0.64).

317

Both of matrix–phase and inclusion–phase plagioclase are Na–rich (An32–43Ab68–57), with no

318

obvious chemical compositional zoning. In most cases, M4 plagioclase has lower anorthite contents

319

(An24–28) than those in other positions. Anorthite in this study has XAn of 0.87–0.94 and XAb of

320

0.13–0.06.

321

Sillimanite typically contains Fe2O3 contents of 0.86–0.93 wt.%.

322 323

5. Metamorphic P–T evolution

324 325

5.1. P–T estimates

326 327

Representative texturally–equilibrated mineral pairs and assemblages are used to estimate P–T

328

conditions through geothermobarometers and average P–T approach. The P–T estimates are

329

tabulated in Table 4.

330

Traditional thermobarometers usually underestimate the peak temperatures of UHT

331

metamorphism as they fail to consider the effects of Fe2+–Mg reset that can occur between mineral

332

pairs during the post–peak cooling process. Therefore, in this contribution, we adopt the garnet–

333

orthopyroxene thermobarometer corrected by Pattison et al. (2003) to estimate the peak P–T

334

conditions of the studied UHT metapelitic granulites. Garnet mantles and high–Al orthopyroxene

335

yielded peak temperatures of 915–1024 °C, with an average value of ~973 °C, approximately

336

representing the temperature of peak metamorphism. The average P–T calculation method of

337

Powell and Holland. (1994) can also be utilized to estimate peak P–T conditions, and the peak

338

assemblage of grt + opx + crd + q + mt + ilm + ru + H2O provides P–T estimates of ~8.8

339

kbar/~980 °C.

340

The

pre–peak

(M1)

assemblage

garnet

(core)–plagioclase–biotite–quartz

(GBPQ)

341

thermobarometry results in P–T conditions of ~4.1 kbar/~690 °C (Holdaway, 2000; Wu et al., 2004).

342

The M3 assemblage gives an average P–T condition of ~7.0 kbar/~790 °C (Powell and Holland,

343

1994). For the post–peak (M4) stage metamorphism, GBPQ thermobarometry yields results of ~6.8

344

kbar/~715 °C (Holdaway, 2000; Wu et al., 2004).

345 346

5.2. Phase equilibria modelling

347 348

The phase equilibria modelling for the UHT granulites sample FY15–49 was carried out using

349

Perple_X (Connolly, 2005) with the internal thermodynamic data set of Holland and Powell, 1998

350

(updated November 2003) in the MnO–Na2O–CaO–K2O–FeO–MgO–Al2O3–SiO2–H2O–TiO2–

351

Fe2O3 (MnNCKFMASHTO) chemical system. The phases considered in the calculation include:

352

garnet (g), orthopyroxene (opx), plagioclase (pl), K–feldspar (ksp), cordierite (crd), biotite (bi),

353

muscovite (ms), sillimanite (sill), kyanite (ky), ilmenite (ilm), magnetite (mt), spinel (sp), quartz

354

(qtz), rutile (ru) and silicate melt (liq). Among these phases, sillimanite, kyanite, rutile and quartz

355

are considered as pure end member phases. The activity–composition (a–x) models for garnet and

356

orthopyroxene are from Holland and Powell. (2001), plagioclase from Newton et al. (1980), K–

357

feldspar from Thompson and Hovis. (1979), cordierite and spinel from Holland and Powell. (1998),

358

biotite from Tajčmanová et al. (2009), muscovite from Coggon and Holland. (2002), ilmenite and

359

magnetite from White et al. (2000), and silicate melt from White et al. (2007).

360 361

5.2.1. Peak and post–peak evolution

362 363

The investigated rock is a residual sample as a consequence of melt loss (SiO2 = 51 wt.%; FeO

364

+ MgO > 20 wt.%). A substantial melt loss is also consistent with the degree of preservation of the

365

peak metamorphic assemblage (White and Powell, 2002). Therefore, the bulk rock composition

366

obtained from XRF analysis is only appropriate for investigating the peak P–T conditions and the

367

post–peak retrograde evolution (White et al., 2004). The X(O) value (representative of ferric iron

368

content) was fixed at 0.28 mol.%, in agreement with values assumed by Li et al., (2014) for similar

369

rocks in this region. The amount of H2O component involved in the calculation was assumed as the

370

loss of ignition (LOI) of XRF analysis.

371

The P–T pseudosection for sample FY15–49 was constructed for the ranges of 2.0–9.0 kbar and

372

700–1100 °C. As is shown in Figure 8a, the solidus is predicted at 750–770 °C for P <5.5 kbar,

373

reflecting the residual bulk chemical composition of this sample. The peak phase assemblage (grt +

374

opx + crd + qtz + mt + ru + liq ± ilm) is represented by the orange fields in Figure 8a and is

375

predicted to be stable at 7–8 kbar and 950–1050 °C, compatible with P–T estimates from the

376

mineral pairs within errors. Rutile– and orthopyroxene–out curves define the lower and upper

377

pressure limits, respectively, whereas biotite disappears just before reaching peak conditions. The

378

proportion of melt predicted at peak conditions is 35–40 vol.%, significantly lower than the amount

379

expected from UHT anatexis of fertile protoliths. This discrepancy is likely due to the residual bulk

380

composition used for calculations.

381

M3 stage is characterized by the formation of opx3–crd3 symplectitic rinds around

382

orthopyroxene and of cordierite + biotite + quartz replacing garnet (see above). In order to consume

383

garnet and melt and produce orthopyroxene, cordierite, biotite and quartz, the post–peak P–T

384

evolution should be characterized by a combination of cooling and decompression (Figs. 8b–g).

385

However, the subsequent M4 stage needs a near–IBC path to form bi4–pl4–qtz intergrowths and

386

consume orthopyroxene, garnet and melt (Figs. 8b–g). Notably, the proposed near–IBC path can

387

also explain the formation of small garnets via reaction (6) at T < 750 °C.

388 389

5.2.2. Prograde evolution

390 391

The reconstruction of a probable bulk chemical composition for the protolith is needed to

392

recover the prograde history of melt–depleted granulites. This approach involves the reintegration

393

of a certain amount of melt to the residual composition and the calculation of phase equilibria for

394

the new protolith composition (Bartoli, 2017). The single–step approach was adopted for the

395

selected sample: the composition of melt in equilibrium with the inferred peak mineral assemblage

396

was calculated at ~7.6 kbar and 1000 °C (SiO2 = 69.50, Al2O3 = 16.67, FeO = 2.55, MgO = 0.77,

397

CaO = 0.96, Na2O = 1.81, K2O = 4.78, H2O = 2.95 wt.%) and an amount of this melt (30%),

398

sufficient to produce a H2O–saturated solidus at ~9 kbar and < 700 °C, was reintegrated.

399

Figure 9 represents the P–T pseudosection for the melt–reintegrated composition. The most

400

evident changes in phase diagram topology are i) the shift of the solidus to lower temperatures at P

401

< 5.5 kbar, ii) the H2O–saturated character of entire solidus curve from 2 to 9 kbar and iii) a

402

reduction of the stability field of orthopyroxene at T < 900 °C (Fig. 9). However, this modelling

403

also fails to predict the pre–peak (M1) mineral phase assemblage. For instance, sillimanite is

404

present only at P > 7 kbar whereas spinel is not predicted to be present. This discrepancy could be

405

related to the presence of chemical–mineralogical microdomains during the prograde evolution of

406

these rocks (Guevara and Caddick, 2016). In this case, the bulk rock composition cannot be

407

representative of the effective bulk composition (EBC) from which such a spinel–bearing mineral

408

assemblage grew.

409

To investigate the possible effect of local variations of the EBC, a T–XMg section based on the

410

melt–reintegrated composition was constructed at 3 kbar, from 650 to 850 °C, with the XMg value

411

ranging from 0 to 1 (Fig. 10). Orthopyroxene–bearing assemblages appear when XMg > 0.2–0.4,

412

whereas garnet is stable for XMg < 0.6–0.7. Assemblages containing spinel or sillimanite are present

413

in the low–XMg side of the diagram (< 0.1). For these low–XMg effective bulk compositions, the

414

cordierite–spinel and cordierite–spinel–sillimanite pre–peak assemblages are predicted to be stable

415

at ~3 kbar, 730–750 °C. The model also indicates that some amounts of garnet could have been

416

already produced during the pre–peak evolution, in agreement with petrographic inferences.

417 418

6. Zircon and monazite U–Th–Pb geochronology

419 420

6.1. Analytical methods

421 422

Conventional magnetic and heavy liquid techniques followed by hand–picking under a

423

binocular microscope were used for separation of zircons from the UHT metapelitic granulites

424

FY15–49 and FY15–51. The morphology and internal structure of the zircons were documented

425

with transmitted and reflected light microphotographs and cathodoluminescene (CL) images.

426

Zircon and monazite U–Th–Pb dating of the samples FY15–49 and FY15–51 were carried out

427

using LA–ICP–MS at the Wuhan Sample Solution Analytical Technology Co., Ltd., Wuhan, China.

428

Detailed operating conditions for the laser ablation system and the ICP–MS instrument and data

429

reduction are described in Zong et al. (2017). Laser sampling was performed using a GeolasPro

430

laser ablation system that consists of a COMPexPro 102 ArF excimer laser (wavelength of 193 nm

431

and maximum energy of 200 mJ) and a MicroLas optical system. Each analysis incorporated a

432

background acquisition of approximately 20–30 s followed by 50 s of data acquisition from the

433

sample.

434

The spot size and frequency of the laser for zircons were set to 32 µm and 10 Hz, respectively.

435

Zircon 91500 and glass NIST610 were used as external standards for U–Pb dating and trace

436

element calibration, respectively. The spot size and frequency of the laser for monazites were set to

437

16 µm and 2 Hz, respectively. Monazite standard 44069 and glass NIST610 were used as external

438

standards for U–Pb dating and trace element calibration, respectively.

439

An Excel–based software ICPMSDataCal was used to perform off–line selection and

440

integration of background and analyzed signals, time–drift correction and quantitative calibration

441

for trace element analyses and U–Pb dating (Liu et al., 2010). Software SQUID 1.0 and ISOPLOT

442

(V.3.0; Ludwig, 1999) were used for data processing.

443 444

6.2. Zircon morphology and U–Pb geochronology

445 446

Representative zircon CL images of UHT metapelitic granulites from the southern Chinese

447

Altai are shown in Figure 11. LA–ICP–MS zircon U–Pb analysis data and age results for the sample

448

FY15–49 are shown in Figure 12a and listed in Table 5. Zircons in this sample are subhedral to

449

anhedral, 60–120 µm in size, and elongate, prismatic, stubby and occasionally round in shape. They

450

usually have core–rim structures, with dark, oscillatory or sector–zoned cores mantled by anhedral

451

overgrowths. Some weekly–zoned or sector–zoned dark cores with bright overgrowths, are also

452

interpreted to be produced during the metamorphism. Thirty–three spots were analyzed on thirty

453

zircon grains from sample FY15–49, with U concentrations of 133–2877 ppm and Th/U ratios of 0–

454

0.87. Zircon rims have relatively low Th/U ratios (≤0.26, mostly less than ~0.1) and lack oscillatory

455

zoning in CL images, suggesting they are metamorphic overgrowths and/or recrystallization

456

features. Among them, six analyses of metamorphic zircon rims show 206Pb/238U ages ranging from

457

266.1 to 297.5 Ma and yield a weighted mean 206Pb/238U age of 281 ± 12 Ma (Fig. 12a). In addition,

458

five analyses of zircon rims have

459

weighted mean

460

phases of metamorphic events in the Chinese Altai orogen. Other grains give

461

428.9–1945.4 Ma, and these are interpreted as xenocrysts. Two

462

320 Ma might reflect mixture ages.

206

Pb/238U ages ranging from 354.5 to 399.0 Ma and produce a

206

Pb/238U age of 393 ± 7.8 Ma (Fig. 12a). These ages indicate the records of two–

206

206

Pb/238U ages of

Pb/238U ages between 360 and

463

CL images of zircons from sample FY15–51 are similar to those from FY15–49 (Fig. 12). LA–

464

ICP–MS zircon U–Pb analysis data and age results are shown in Fig. 12b and listed in Table 6.

465

Forty–three spots were analyzed on thirty–seven zircon grains for sample FY15–51, with U

466

concentrations of 87–1210 ppm and Th/U ratios of 0.01–0.83. Zircon rims have relatively low Th/U

467

ratios (≤ 0.05) and lack oscillatory zoning in CL images, suggesting they are metamorphic

468

overgrowths and/or recrystallization features. Among them, twelve analyses show

469

ranging from 261.4 to 300.6 Ma and yield a weighted mean

470

12b). Additionally, nine analyses of zircon rims have

471

Ma and yield a weighted mean

472

two–phases of metamorphic events in the southern Chinese Altai. Igneous zircons give

473

ages ranging from 441.6 Ma to 1250.8 Ma. Other obtained 206Pb/238U ages between 360 and 320 Ma

474

are interpreted as mixture ages.

206

Pb/238U ages

206

Pb/238U age of 282 ± 11 Ma (Fig.

206

Pb/238U ages ranging from 364.1 to 397.6

206

Pb/238U age of 387 ± 6.7 Ma (Fig. 12b). These ages also record 206

Pb/238U

475 476

6.3. Monazite morphology and U–Th–Pb geochronology

477 478

Most monazites in the sample FY15–51 are 50–250 µm in size，round, stubby or irregular in

479

shape, and weekly core–rim zoned

(Figs. 13c–f). They generally occur as inclusions preserved in

480

garnet or cordierite (Figs. 13a and b). Eighteen available analyses were conducted on nine

481

monazites of sample FY15–51. Their

482

(Table 7). The most concordant 18 data give a weighted mean

483

(Fig. 14a). This cluster includes monazite from different textural positions, and thus it does not

484

reveal statistically distinguishable ages. The weighted mean U–Pb concordia age is 289.45 ± 1.4 Ma

206

Pb/238U ages show a range from 293.3 Ma to 271.9 Ma 208

Pb/232Th age of 287.9 ± 2.1 Ma

208

Pb/232Th age within the analytical

485

(Fig. 14b), which is in agreement with the weighted mean

486

errors. The lower intercept age (277 ± 30 Ma) is slightly younger than the

487

absence of a microstructural control on the monazite ages may suggest a complete resetting of the

488

U–Th–Pb isotope system that probably affected all the monazite crystals (Langone et al., 2010).

489

Accordingly, the age of 287.9 ± 2.1 Ma is regarded as an appropriately estimated metamorphic age

490

for this sample.

208

Pb/232Th age. The

491 492

7. Discussion

493 494

7.1. P–T path and timing constraints

495 496

UHT metamorphism has been documented in two localities in the southern Chinese Altai (Figs.

497

1 and 2): Kalasu area (Tong et al., 2013, 2014), and Wuqiagou area (Li et al., 2010, 2014; Yang et

498

al., 2015b). Both of clockwise and anticlockwise P–T paths were defined for the Altai UHT

499

granulites (Fig. 15), as discussed above. In this study, P–T conditions of four different stages define

500

an anticlockwise P–T path with peak P–T conditions of ∼8.0 kbar/~980 °C, which is roughly

501

consistent with that of Tong et al. (2014). P–T path proposed by Tong et al. (2014) involved a short

502

post–peak P–T fragment from UHT conditions to ~870 °C at 8–9 kbar. Our UHT granulite samples,

503

however, recorded an integrated P–T path with a post–peak decompression and a subsequent near–

504

IBC process from ~900 °C (Fig. 15). It is important to note that the pressure and temperature

505

constraints obtained from classic geothermobarometers and average P–T approach would suggest

506

the presence of near–IBC path characterized by a minimal decompression (Fig. 15). However, this

507

P–T evolution is not consistent with the formation of coronas and thermodynamic calculations. It

508

follows that the P–T path constructed from phase equilibria modelling is more reliable (see above).

509

A combination of zircon and monazite chronological data generally allows for a higher

510

temporal resolution of high–grade metamorphism than when each method is applied separately (e.g.,

511

Wu et al., 2014). The LA–ICP–MS zircon U–Pb age data obtained from the Altai UHT granulites

512

indicate the existence of at least two metamorphic events. Weighted mean

513

7.8 and 387 ± 6.7 Ma yielded from metamorphic zircons are consistent with the timing proposed for

514

the regionally extensive Devonian metamorphism (390–365 Ma; Zhuang, 1994; Hu et al., 2002;

515

Windley et al., 2002; Wei et al., 2007; Zheng et al., 2007; Jiang et al., 2010; Broussolle et al., 2018).

516

Considering M1 metamorphic temperature conditions (700–750 °C) are roughly in agreement with

517

those (650–700 °C) of amphibolite and paragneiss in the southern Chinese Altai (Jiang et al., 2010),

518

M1 stage may be reasonably inferred to be linked with the Devonian tectono–metamorphic event.

206

Pb/238U ages of 393 ±

519

Furthermore, weighted mean 206Pb/238U ages of ∼280 Ma produced from metamorphic zircons

520

are in agreement with the Permian tectono–metamorphic event (295–260 Ma; Chen et al., 2006;

521

Briggs et al., 2007; Zheng et al., 2007; Wang et al., 2009b; Li et al., 2014; Tong et al., 2014). Our

522

monazite U–Th–Pb dating results (287.9 ± 2.1 Ma) further support the existence of the Permian–age

523

metamorphic reworking event. We interpret these ages as the timing of the Altai UHT

524

metamorphism. Our monazite age dating results, however, did not document the Devonian

525

metamorphic event, which is probably because monazite recrystallized and reset during UHT

526

metamorphism (Wu et al., 2014; Morrissey et al., 2016).

527

Granitoids are extensively cropping out at Wuqiagou area and also documented two prominent

528

thermal–magmatic events of 393.5 ± 4.5 and 293.5 ± 6.0 Ma (Zhang et al., 2015a). Structurally, the

529

southern Chinese Altai documented two episodes of deformation D1 and D2, among which the

530

early sub–horizontal fabric S1 was modified to different degrees by the late near–vertical

531

deformation D2 in the Devonian (Jiang et al., 2015). Locally, the originally S1 and D2 fabrics were

532

reworked by the Permian upright close and asymmetric F3 folding (Broussolle et al., 2018). Our

533

zircon U–Pb ages also provide a solid support that the Altai UHT granulites experienced two

534

prominent tectono–metamorphic cycles: the first Devonian (~390 Ma) metamorphic and the second

535

Permian (~280 Ma) reworking process.

536 537

7.2. Tectonic Implications

538 539

A long–lived single arc north–dipping subduction and multiple arc subduction of the

540

Kazakhstan–Junggar plate have been proposed to explain the continental growth of Altai orogen in

541

Phanerozoic (e.g., Briggs et al., 2007; Windley et al., 2007). Two tectono–metamorphic cycles with

542

time interval of ~100 Ma are recognized and considered to have affected the whole edifice of the

543

Chinese Altai (Zhuang et al., 1994; Wei et al., 2007; Broussolle et al., 2018), as described above.

544

The first tectono–metamorphic cycle is extensive in the whole Chinese Altai Belt and took place in

545

the middle Devonian (Wei et al., 2007; Jiang et al., 2010; Broussolle et al., 2018). The Devonian

546

orogenic architecture was subsequently reworked by the Permian HT–UHT metamorphism on its

547

southern margin (Wang et al., 2009b, 2014; Li et al., 2014; Tong et al., 2014; Liu and Tong, 2015).

548

Particularly, previous zircon U–Pb and monazite U–Th–Pb dating results of Altai high

549

temperature gneisses and granulites showed metamorphic ages of 293–260 Ma (Chen et al., 2006;

550

Briggs et al., 2007; Zheng et al., 2007; Wang et al., 2009b), indicating that the Chinese Altai

551

orogenic belt experienced a significant HT–UHT reworking event in the late Paleozoic (Xiao et al.,

552

2008). In terms of the Permian HT–UHT metamorphism, researchers have proposed various genetic

553

models, such as Paleo–Asian oceanic crustal subduction (Li et al., 2004; Chen et al., 2006; He et al.,

554

2018), collision of the Junggar arc with the Chinese Altai terrane (Li et al., 2010; Broussolle et al.,

555

2018), slab break–off (Li et al., 2014; Yang et al., 2015b), ridge–subduction and the development of

556

slab–window (Windley and Xiao, 2018), and the thermal pulse of Tarim mantle plume activities

557

(Tong et al., 2013, 2014; Wang et al., 2014; Liu and Tong, 2015; Yang et al., 2015a). However,

558

until recently, the tectonic nature of the southern Chinese Altai is still a controversial topic.

559

The P–T path may be diagnostic of a particular tectonic environment. In particular, the

560

anticlockwise P–T path coupled with UHT peak conditions and a post–peak near–IBC process

561

generally reflects a tectonic evolutionary history involving initial crustal thickening in areas of

562

voluminous magmatic accretion (Fig. 15) (Harley, 1989; Sandiford and Powell, 1991). In addition,

563

the occurrence of a decompression segment is consistent with the extensional thinning of a

564

thickened crust (Harley, 1989). The fact that crustal thickening characterized by a subsequent near–

565

IBC is normally accompanied by underplating/accretion of deep–derived magma has been

566

discussed in many studies (e.g., Brandt et al., 2003; Clark et al., 2014; Jiao et al., 2015). In the

567

southern Chinese Altai, Permian granites have A– or I/A–type characteristics (Wang et al., 2010;

568

Shen et al., 2013), with a few high–temperature S–type granites (Zhou et al., 2007). These

569

granitoids and voluminous contemporary mafic intrusions show a bimodal magmatic association

570

(e.g., Wang et al., 2010; Shen et al., 2013), consistent with a post–orogenic extensional tectonic

571

setting (Shen et al., 2013, Wang et al., 2014). Accordingly, the models involving only a simple

572

subduction process or an arc–continent collision seem to be impossible.

573

The slab break–off was proposed by Li et al. (2014) and Yang et al. (2015b) to explain the

574

UHT metamorphism at Wuqiagou area, which we also considered unlikely. Freeburn et al. (2017)

575

proposed that in most cases slab break–off occurs too deep to trigger melting and cause thermal

576

perturbation within the overriding plate. On the other side, mantle wedge might be unable to

577

provide enough free space for voluminous hot asthenosphere to flow from beneath the slab and fill

578

the wedge above the slab (Niu et al., 2017). Collectively, the model of slab detachment might be

579

unable to explain the regionally extensive magmatic activities, crustal melting and the late orogenic

580

extension in the Chinese Altai orogenic belt.

581

A regional significant event that is broadly coeval with this HT–UHT metamorphism was the

582

formation of the Permian–aged Large Igneous Province (LIP) in northwestern China (Zhang et al.,

583

2012; Liu et al., 2014; Xu et al., 2014). The Permian Tarim LIP covering an area of more than 250

584

000 km2 (Tian et al., 2010; Yang et al., 2006) has recently become a focus of research (Xu et al.,

585

2014 and references therein). Some recent studies have suggested that the coeval Permian mafic

586

magmatic activities were associated with the rapid ascent of plume–related magma beneath the

587

lithosphere (e.g., Chen and Han, 2006; Zhang et al., 2012; Shen et al., 2013; Yang et al., 2015a).

588

Considering the time consistence with the Tarim mantle plume activity (~275 Ma; Zhang et al.,

589

2012), we proposed that the Altai HT–UHT metamorphism was most likely linked to the

590

underplating and heating of mantle–derived mafic magma as a result of the Tarim mantle plume

591

activity. Although we could not exclude the possibility of ridge subduction, the Tarim mantle plume

592

activity seems to be a better genetic model.

593

A possibly textural evidence for the existence of the Tarim mantle plume activity in the

594

southern Chinese Altai is the sinistral strike–slip motion of the Erqis fault belt (290–280 Ma;

595

Laurent–Charvet et al., 2003; Briggs et al., 2007; Zhang et al., 2012). Both of the Altai UHT

596

granulites and low–pressure metapelitic gneisses occurred along the large Erqis fault belt. The

597

movement of Erqis fault belt was ascribed to be at least partially influenced by the Tarim mantle

598

plume activity (Zhang et al., 2012). This probably resulted in the emplacement of voluminous A–

599

type granites and post–orogenic mafic–ultramafic intrusions during the same time. Accordingly, the

600

high heat flow necessary for the HT–UHT metamorphism of the southern Chinese Altai was

601

provided by coeval mafic intrusions which were probably generated by the Tarim mantle plume

602

activity.

603 604

8. Conclusions

605 606

Through detailed petrographic observations and P–T estimates for the UHT metapelitic

607

granulites from the southern Chinese Altai, in combination with both zircon and monazite U–Th–Pb

608

geochronology, some suggestions can be derived as follows:

609

(1) Petrography, mineral compositions and metamorphic P–T estimates for the Altai Permian

610

metapelitic granulites from Wuqiagou area suggest an anticlockwise P–T path characterized by

611

UHT peak conditions (~980 °C), and a post–peak decompression and a subsequent near–IBC

612

processes.

613

(2) The LA–ICP–MS zircon U–Pb dating results indicate the existence of at least two separated

614

high–grade metamorphic events. The M1 stage might occur in the middle Devonian (~390 Ma).

615

And then, the Devonian metamorphic terranes were locally overprinted by the Permian–age

616

UHT metamorphism (∼280 Ma). Monazite U–Th–Pb age dating results for our samples (287.9

617 618 619

± 2.1 Ma) provide a further constraint on the timing of the UHT reworking event. (3) The Altai Permian UHT reworking event was most likely associated with the underplating and heating of deep–derived mafic magma as a result of the Tarim mantle plume activity.

620 621

Acknowledgements

622 623

This study has been supported by the Strategic Priority Research Program (B) of the Chinese

624

Academy of Sciences (XDB18030601), a One Hundred Talents Project of Shaanxi Province granted

625

to L. Tong, and by SIR RBSI14Y7PF grant by Italian Ministry of Education, University, Research

626

to O. Bartoli. We are really grateful to China Scholarship Council for its financial support during a

627

visit of Zhao Liu to Università di Padova, Italy. The Electron Microprobe analysis for mineral

628

composition was finished with help of Ms L.L. Chen at State Key Lab of Isotope Geochemistry,

629

Guangzhou Institute of Geochemistry. LA–ICP–MS zircon and monazite analysis were completed

630

with help of Mr Wei Gao at Wuhan Sample Solution Analytical Technology Co., Ltd., Wuhan,

631

China. We very much appreciate the journal editor Prof. T. Tsunogae and two anonymous reviewers

632

for their helpful and constructive comments on the early version of this paper.

633 634

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intrusive mass in Wuqiagou on the southern margin of Altay Orogenic Belt and its geological

895

significance. Northwestern Geology 48, 127–139 (in Chinese with English abstract).

896

Zhang, Z., Dong, X., Xiang, H., Ding, H., He, Z., Liou, J.G., 2015b. Reworking of the Gangdese

897

magmatic arc, southeastern Tibet: post–collisional metamorphism and anatexis. Journal of

898

Metamorphic Geology 33: 1–21.

899

Zhao, L., Guo, F., Fan, W.M., Li, C.W., Qin, X.F., Li, H.X., 2010. Origin of the granulite enclaves

900

in Indo–Sinian peraluminous granites, South China and its implication for crustal anatexis.

901

Lithos 150: 209–226.

902

Zheng, C.Q., Xu, X.C., Kato, T., Enami, M., 2007. Permian CHIME ages of monazites for the

903

kyanite–sillimanite type metamorphic belt in Chonghuer area, Altai, Xinjiang and their

904

geological implications. Geological Journal of China University 13, 566–573 (in Chinese with

905

English abstract).

906

Zhou, G., Zhang, Z.C., Luo, S.B., He, B., Wang, X., Ying, L.J., Zhao, H., Li, A.H., He, Y.K., 2007.

907

Confirmation of high–temperature strongly peraluminous Mayin’ebo granites in the margin of

908

Altay, Xinjiang: age, geochemistry and tectonic implications. Acta Petrologica Sinica 23, 1909–

909

1920 (in Chinese with English abstract).

910

Zhuang, Y.X., 1994. The PTSt evolution of metamorphism and development mechanism of the

911

thermal–structural–gneiss domes in the Chinese Altaides. Acta Geologica Sinica 68, 35–47 (in

912

Chinese with English abstract).

913

Zong, K.Q., Klemd, R., Yuan, Y., He, Z.Y., Guo, J.L., Shi, X.L., Liu, Y.S., Hu, Z.C., Zhang, Z.M.,

914

2017. The assembly of Rodinia: The correlation of early Neoproterozoic (ca. 900 Ma) high–

915

grade metamorphism and continental arc formation in the southern Beishan Orogen, southern

916

Central Asian Orogenic Belt (CAOB). Precambrian Research 290, 32–48.

917 918

FIGURE CAPTIONS

919 920

Fig. 1. A simplified metamorphic geological map of the Chinese Altai (modified after Wei et al.,

921

2007).

922 923

Fig. 2. A geological sketch map of the Fuyun area and sampling location of the Altai granulites

924

(modified after Li et al., 2004).

925

★: Sample location: O2–3: Middle to Late Ordovician Habahe Group; Sk1: Silurian Kulumuti Group;

926

D1k: Early Devonian Kangbutibao Group; D2a: Middle Devonian Aletai Group ； C3k: Late

927

Carboniferous Kala–Erqis Group; J3: Late Jurassic Shishugou Group; Cz: Cenozoic Group; ψ4:

928

Variscan mafic and ultramafic rocks; γ4: Variscan granitic rocks.

929 930

Fig. 3. The field photograph of the UHT metapelitic granulites from the Wuqiagou area in the

931

southern Chinese Altai. The marker pen for scale is about 10 cm long.

932 933

Fig. 4. Photomicrographs and back–scattered electron (BSE) images showing the pre–peak and

934

peak mineral assemblages and textures in the UHT metapelitic granulites from Wuqiagou area in

935

southern Chinese Altai.

936

(a), spinel–cordierite–magnetite–anorthite associations in the matrix (FY15–49); (b), a garnet

937

porphyroblast with inclusions of spinel, cordierite, biotite, plagioclase and magnetite (FY15–49);

938

(c), a cordierite porphyroblast with spinel, cordierite, biotite, sillimanite and magnetite inclusions

939

(FY15–51); (d), a peak rutile grain preserved in retrograde quartz (FY15–49); (e), plagioclase and

940

quartz inclusions preserved in the orthopyroxene porphyroblast (FY15–49); (f), inclusion–phase

941

biotite, cordierite and magnetite contained in the orthopyroxene porphyroblast (FY15–51); (g), melt

942

pseudomorphs occurring among plagioclase. The outlines of quartz pools and films are indicative of

943

crystallization from melts (FY15–49); (h), MI clusters preserved in peritectic garnet (FY15–49); (i)

944

and (j), BSE images of selected nanogranites preserved in garnets.

945

Mineral abbreviations: grt, garnet; opx, orthopyroxene; sp, spinel; crd, cordierite; bi, biotite; sill,

946

sillimanite; pl, plagioclase; an, anorthite; qtz, quartz; mt, magnetite; ilm, ilmenite; ru, rutile; apt,

947

apatite; monz, monazite; zr, zircon; MI, melt inclusions.

948 949

Fig. 5. Photomicrographs and back–scattered electron (BSE) images illustrating reaction textures

950

resulting from the breakdown of garnet and orthopyroxene.

951

(a), opx2 rimed by opx–crd symplectitic rinds (FY15–51); (b), a BSE image illustrating the opx–crd

952

symplectite resorbing opx2 (FY15–51); (c), ‘spongy’ appearance around garnet consisted of

953

cordierite moats (crd3) and vermiform quartz (FY15–49); (d), the enlarged BSE image of black box

954

in (c) showing crd–qtz symplectite replacing garnet (FY15–49); (e) a photograph showing the

955

regrowth of the late garnet around orthopyroxene (FY15–51); (f), the intergrowths of bi–pl–qtz

956

destructing the early orthopyroxene (FY15–51); (g), the intergrowths of bi–pl–qtz bypassing the

957

garnet porphyroblast, suggesting their retrograde origin (FY15–49); (h), rutile needles preserved in

958

the garnet (FY15–49).

959

Mineral abbreviations see Figure 4.

960 961

Fig. 6. (SiO2)–A(Al2O3 + Fe2O3)–FM(FeO + MgO) projection from plagioclase, K–feldspar and

962

biotite for bulk rock compositions of the samples.

963 964

Fig. 7. (a) and (b), compositional variations of garnets from the Altai UHT granulites; (c)–(f),

965

compositional zoning profiles across garnet porphyroblasts.

966 967

Fig. 8. (a), P–T pseudosection for the UHT granulite sample FY15–49, calculated in

968

MnNCKFMASHTO system and considering the residual bulk rock composition. Orange fields

969

reflect the predicted mineral assemblages. (b–g), mineral and melt proportions (in vol.%). Green

970

arrow represents the probable post–peak retrograde evolution.

971 972

Fig. 9. P–T pseudosection for the UHT granulite sample FY15–49, calculated in

973

MnNCKFMASHTO system and considering the melt–reintegrated composition.

974 975

Fig. 10. T–XMg pseudosection constructed in MnNCKFMASHT system at 3 kbar and showing

976

different mineral assemblages due to heterogeneous effective bulk compositions.

977 978

Fig. 11. Cathodoluminescence (CL) images and 206Pb/238U ages of the zircon grains separated from

979

the Altai UHT granulite samples FY15–49 and FY15–51, respectively. The circles stand for

980

analytical spots and the neighboring white numbers are the respective 206Pb/238U ages.

981 982

Fig. 12. U–Pb concordant age diagrams showing the LA–ICP–MS zircon age results for FY15–49

983

and FY15–51, respectively.

984 985

Fig. 13. Representative photomicrographs and BSE images of analyzed monazite grains for sample

986

FY15–51 and locations of the analytical spots. The individual apparent

987

labeled.

988

(a) and (b), monazite grains within matrix and cordierite porphyroblasts; (c), an enlarged BSE

989

image of the monazite grain in (a); (d), a monazite inclusion preserved in garnet; (e) and (f),

990

enlarged BSE images of monazite grains in (b).

991

208

Pb/232Th age is also

208

Pb/232Th age for sample FY15–51; (b), concordia diagrams of

992

Fig. 14. (a), the weighted mean

993

monazite LA–ICP–MS U–Pb analytical results for sample FY15–51.

994 995

Fig. 15. Suggested P–T path of the UHT granulites from the southern Chinese Altai (after Wei et

996

al., 2007). Biotite–dehydration reactions in the KFMASHTO system from White et al. (2002) are

997

marked in this figure. Also shown are P–T estimates and P–T paths suggested by Wang et al.

998

(2009b, 2014), Li et al. (2014), Tong et al. (2014) and Yang et al. (2015b) for the HT–UHT rocks

999

in this region.

1000 1001

Table captions:

1002 1003

Table 1. Bulk rock compositions of the UHT granulites at Wuqiagou area.

1004 1005

Table 2. Representative garnet and orthopyroxene compositions of the Altai UHT granulites.

1006 1007

Table 3. Representative spinel, cordierite, biotite and plagioclase compositions of the Altai UHT

1008

granulites.

1009 1010

Table 4. P–T estimates for four–stage mineral assemblages in the UHT metapelitic granulites from

1011

the southern Chinese Altai.

1012

Abbreviations: H00 (Holdaway, 2000); PH94 (Powell and Holland, 1994); P03 (Pattison et al.,

1013

2003); W04 (Wu et al., 2004).

1014 1015

Table 5. LA–ICP–MS U–Th–Pb analysis results for zircons from the Altai UHT granulite sample

1016

FY15–49.

1017 1018

Table 6. LA–ICP–MS U–Th–Pb analysis results for zircons from the Altai UHT granulite sample

1019

FY15–51.

1020 1021

Table 7. LA–ICP–MS monazite U–Th–Pb isotopic analyses for the Altai UHT granulite sample

1022

FY15–51.

1023

Table 1 Sample

SiO2

FY15-49 51.42 FY15-51 56.49

Al2O3

Fe2O3T MgO

CaO

MnO

Na2O

K2O

TiO2

P2O5

Cr2O3

LOI

Total

19.31 18.72

15.24 11.56

0.65 0.94

0.39 0.20

0.64 0.96

1.71 1.90

1.48 1.24

0.04 0.08

0.06 0.04

1.12 1.32

100.58 100.35

8.52 6.87

Table 2

core SiO2 TiO2 Al2O3 Cr2O3 FeO MnO MgO CaO Na2O K2O Total Cation(O) Si Ti Al Cr Fe3+ Fe2+ Mn Mg Ca Na K Sum

37.70 0.02 20.32 0.12 32.04 2.97 5.50 1.21 0.07 0.00 99.95 12 2.995 0.001 1.903 0.008 0.108 2.020 0.200 0.651 0.103 0.011 0.000 8.000

Alm Py Sp Grs Mg#

0.68 0.22 0.07 0.03 24

core

garnet mantle mantle

rim

38.37 38.86 39.05 38.56 0.00 0.04 0.06 0.03 21.40 21.68 21.71 21.29 0.07 0.06 0.00 0.04 30.86 29.41 28.44 31.33 2.04 1.93 1.83 2.71 6.83 7.68 8.30 5.75 1.25 1.12 1.20 1.13 0.01 0.02 0.05 0.00 0.01 0.01 0.00 0.00 100.84 100.81 100.64 100.84 12 12 12 12 2.988 3.006 3.010 3.022 0.000 0.002 0.003 0.002 1.964 1.977 1.973 1.967 0.004 0.004 0.000 0.002 0.058 0.007 0.007 0.000 1.951 1.896 1.827 2.053 0.135 0.126 0.119 0.180 0.793 0.885 0.954 0.671 0.104 0.093 0.099 0.095 0.002 0.003 0.007 0.000 0.001 0.001 0.000 0.000 8.000 8.000 7.999 7.992 0.65 0.27 0.05 0.03 29

0.63 0.30 0.04 0.03 32

0.61 0.32 0.04 0.03 34

0.68 0.22 0.06 0.03 25

orthopyroxene rim rim symplectite symplectite

rim

grt4

core

core

core

37.82 0.05 20.70 0.05 31.50 3.14 4.58 1.66 0.01 0.01 99.52 12 3.026 0.003 1.952 0.003 0.000 2.107 0.213 0.546 0.142 0.002 0.001 7.995

38.20 0.00 20.82 0.00 30.99 3.20 5.84 1.17 0.06 0.01 100.29 12 3.012 0.000 1.935 0.000 0.052 1.991 0.214 0.686 0.099 0.010 0.001 8.000

51.73 0.07 8.71 0.10 21.66 0.61 16.05 0.31 0.80 0.01 100.05 6 1.907 0.002 0.379 0.003 0.000 0.668 0.019 0.882 0.012 0.057 0.000 3.929

50.73 0.01 8.14 0.12 22.06 0.71 16.27 0.33 0.67 0.00 99.04 6 1.899 0.000 0.359 0.004 0.000 0.690 0.023 0.908 0.013 0.049 0.000 3.945

52.90 0.12 8.07 0.12 20.01 0.46 17.16 0.24 0.65 0.01 99.74 6 1.936 0.003 0.348 0.003 0.000 0.612 0.014 0.936 0.009 0.046 0.000 3.907

0.70 0.18 0.07 0.05 21

0.67 0.23 0.07 0.03 26

0.43 0.57 0.18

0.40 0.60 0.17

0.40 0.60 0.13

57

60

60

En 0.43 Fs 0.57 AlⅣ 0.19 57

49.58 49.54 0.00 0.07 5.68 5.85 0.07 0.04 24.24 24.15 0.54 0.50 19.88 19.89 0.06 0.06 0.02 0.00 0.00 0.02 100.07 100.12 6 6 1.857 1.854 0.000 0.002 0.251 0.258 0.002 0.001 0.034 0.029 0.725 0.727 0.017 0.016 1.110 1.109 0.002 0.002 0.001 0.000 0.000 0.001 3.999 3.999

55.81 0.12 2.25 0.03 21.38 0.82 19.07 0.14 0.05 0.34 100.01 6 2.052 0.003 0.098 0.001 0.000 0.657 0.026 1.045 0.006 0.004 0.016 3.908

56.02 0.03 2.00 0.08 21.20 0.80 20.15 0.15 0.06 0.00 100.49 6 2.045 0.001 0.086 0.002 0.000 0.647 0.025 1.096 0.006 0.004 0.000 3.912

0.40 0.60 0.13

0.39 0.61 0.05

0.37 0.63 0.04

60

61

63

Table 3 spinel matrix matrix sp in grt

cordierite matrix crd in sp crd in grt

SiO2 TiO2 Al2O3 Cr2O3 FeO MnO MgO CaO Na2O K2O Total Cation(O) Si Ti Al Cr Fe3+ Fe2+ Mn Mg Ca Na K Sum

0.45 0.08 57.44 0.26 34.77 0.24 5.80 0.00 0.03 0.01 99.08 4 0.013 0.002 1.909 0.006 0.059 0.761 0.006 0.244 0.000 0.002 0.000 3.002

0.02 0.00 59.30 0.42 28.71 0.11 7.96 0.00 0.06 0.01 96.59 4 0.001 0.000 1.972 0.009 0.021 0.656 0.003 0.335 0.000 0.003 0.000 3.000

49.56 0.00 33.26 0.73 5.38 0.10 9.75 0.02 0.14 0.00 98.94 18 5.010 0.000 3.964 0.058 0.000 0.455 0.009 1.469 0.002 0.027 0.000 10.994

48.54 0.00 33.49 0.50 5.58 0.15 10.12 0.05 0.24 0.00 98.67 18 4.924 0.000 4.005 0.040 0.095 0.379 0.013 1.530 0.005 0.028 0.000 11.019

Mg# An Ab

24.000 34.000 34.000

76.000

80.000

0.04 0.00 57.39 0.39 33.80 0.23 5.60 0.00 0.10 0.00 97.55 4 0.001 0.000 1.972 0.009 0.021 0.656 0.003 0.335 0.000 0.003 0.000 3.000

crd3

biotite rim bi in grt bi in crd

crd3

core

50.12 0.00 32.75 0.14 3.80 0.06 11.27 0.02 0.15 0.00 98.31 18 5.050 0.000 3.890 0.011 0.029 0.292 0.005 1.692 0.002 0.029 0.000 11.000

49.65 49.16 0.01 0.00 31.70 31.69 0.06 0.02 4.61 4.70 0.14 0.05 10.33 10.48 0.00 0.02 0.22 0.15 0.20 0.00 96.92 96.27 18 18 5.098 5.080 0.001 0.000 3.837 3.861 0.005 0.002 0.030 0.007 0.366 0.400 0.012 0.004 1.581 1.614 0.000 0.002 0.044 0.030 0.026 0.000 11.000 11.000

37.52 3.18 15.69 0.05 14.21 0.05 14.65 0.00 0.45 8.95 94.75 11 2.797 0.178 1.379 0.003 0.000 0.886 0.003 1.628 0.000 0.065 0.851 7.792

37.40 3.35 16.00 0.08 14.54 0.09 14.54 0.02 0.22 8.89 95.13 11 2.779 0.187 1.402 0.005 0.000 0.904 0.006 1.610 0.002 0.032 0.843 7.768

85.000

81.000

65.000

64.000 67.000 57.000

80

37.82 1.51 17.11 0.16 14.45 0.04 14.81 0.06 0.12 9.01 95.09 11 2.793 0.084 1.489 0.009 0.073 0.820 0.003 1.630 0.005 0.017 0.849 7.772

35.65 4.99 16.34 0.12 15.87 0.12 11.69 0.00 0.36 9.25 94.39 11 2.707 0.285 1.463 0.007 0.000 1.008 0.008 1.323 0.000 0.053 0.896 7.750

plagioclase matrix pl in grt pl in opx

pl4

59.80 0.00 24.28 0.00 0.06 0.01 0.00 7.05 8.04 0.07 99.31 8 2.687 0.000 1.286 0.000 0.002 0.000 0.000 0.000 0.339 0.701 0.004 5.019

57.01 0.00 26.47 0.02 0.25 0.02 0.00 9.15 6.55 0.03 99.50 8 2.572 0.000 1.408 0.001 0.009 0.000 0.001 0.000 0.442 0.573 0.002 5.008

59.43 0.00 24.56 0.02 0.17 0.02 0.00 7.51 7.67 0.03 99.41 8 2.670 0.000 1.301 0.001 0.006 0.000 0.001 0.000 0.362 0.668 0.002 5.011

60.75 0.00 23.58 0.02 0.17 0.00 0.01 6.14 8.81 0.05 99.53 8 2.721 0.000 1.245 0.001 0.006 0.000 0.000 0.001 0.295 0.765 0.003 5.037

0.32 0.67

0.43 0.56

0.35 0.65

0.28 0.72

Table 4 M1

T = ~690 °C

H00

M2

P03 at P = 7 kbar

Average PH94

Modle 1 Modle 2 T = 915 °C T = 975 °C T = 950 °C T = 999 °C T = 940 °C T = 1000 °C T = 974 °C T = 1024 °C — — P = ~8.8 kbar, T = ~980 °C

Modle 3 T = 979 °C T = 1002 °C T = 1004 °C T = 1027 °C — cor = 0.54, sigfit = 1.73

Modle 4 Average T = 914 °C T = 946 °C T = 951 °C T = 976 °C T = 939 °C T = 971 °C T = 975 °C T = 1000 °C — T = 973 °C When a(H2O) = 0.3, a(CO2) = 0.7

PH94

P = ~7.0 kbar, T = ~790 °C

cor = 0.41

sigfit = 1.23

at P = 8 kbar

M3

P = ~4.10 kbar

W04

T = ~715 °C H00 P = ~6.85 kbar W04 M4 Abbreviations: H00 (Holdaway, 2000); PH94 (Powell and Holland, 1994); P03 (Pattison et al., 2003); W04 (Wu et al., 2004).

Table 5 U

Th

Th/U

FY15-49-01

757

7

0.01

FY15-49-02

1136

7

FY15-49-03

697

17

FY15-49-04

1096

5

FY15-49-05

925

23

FY15-49-06

785

7

FY15-49-07

680

FY15-49-08

471

FY15-49-09

Spot

206

Pb/238U

1σ

207

Pb*/206Pb*

1σ

207

Pb*/U235

1σ

206

Pb/238U(Ma) 1σ

207

Pb/206Pb(Ma)

1σ

207

Pb/235U(Ma)

1σ

0.042

0.00045

0.052

0.00134

0.304

0.00762

266.1

2.80

301.9

62.03

269.7

5.93

0.01

0.043

0.00038

0.055

0.00120

0.330

0.00717

274.5

2.37

405.6

43.52

289.5

5.48

0.02

0.045

0.00049

0.058

0.00243

0.361

0.01713

281.0

3.02

542.6

90.73

313.2

12.78

0.00

0.046

0.00049

0.053

0.00117

0.336

0.00750

287.9

3.02

342.7

50.00

294.2

5.70

0.03

0.047

0.00071

0.053

0.00153

0.350

0.01214

296.9

4.35

350.1

64.81

305.1

9.13

0.01

0.047

0.00058

0.055

0.00126

0.354

0.00832

297.5

3.58

394.5

56.48

307.8

6.24

25

0.04

0.057

0.00116

0.053

0.00114

0.418

0.01210

354.5

7.07

350.1

50.00

354.7

8.66

20

0.04

0.055

0.00102

0.055

0.00135

0.417

0.01347

345.7

6.26

390.8

55.55

354.2

9.65

507

21

0.04

0.063

0.00136

0.057

0.00251

0.494

0.02435

390.9

8.27

487.1

102.77

407.7

16.55

FY15-49-10

437

26

0.06

0.063

0.00104

0.050

0.00134

0.436

0.01394

393.1

6.33

183.4

67.58

367.3

9.86

FY15-49-11

739

6

0.01

0.063

0.00106

0.057

0.00264

0.497

0.02190

393.7

6.45

494.5

101.84

409.7

14.85

FY15-49-12

536

138

0.26

0.063

0.00134

0.059

0.00178

0.514

0.01640

396.2

8.09

576.0

66.66

421.3

11.00

FY15-49-13

944

187

0.20

0.064

0.00114

0.052

0.00120

0.466

0.01319

399.0

6.92

301.9

51.85

388.2

9.14

FY15-49-14

2877

366

0.13

0.158

0.00128

0.066

0.00189

1.461

0.03923

943.5

7.10

820.4

59.26

914.5

16.19

FY15-49-15

932

190

0.20

0.082

0.00070

0.057

0.00153

0.650

0.01694

509.7

4.16

487.1

59.25

508.3

10.43

FY15-49-16

558

239

0.43

0.084

0.00120

0.057

0.00138

0.670

0.01792

519.7

7.14

505.6

56.48

520.8

10.90

FY15-49-17

220

83

0.37

0.089

0.00169

0.059

0.00168

0.726

0.02320

549.0

10.02

572.3

61.87

553.9

13.65

FY15-49-18

565

443

0.78

0.088

0.00119

0.060

0.00124

0.731

0.01709

544.7

7.05

598.2

46.29

557.3

10.03

FY15-49-19

520

98

0.19

0.078

0.00101

0.058

0.00136

0.631

0.01677

481.8

6.06

542.6

47.22

496.5

10.44

FY15-49-20

646

139

0.22

0.080

0.00130

0.058

0.00120

0.634

0.01464

496.2

7.77

522.3

50.92

498.6

9.10

FY15-49-21

317

56

0.18

0.070

0.00085

0.055

0.00166

0.533

0.01649

433.9

5.15

433.4

66.66

434.0

10.92

FY15-49-22

431

27

0.06

0.071

0.00138

0.055

0.00147

0.540

0.01616

442.5

8.29

433.4

59.25

438.6

10.65

FY15-49-23

203

151

0.74

0.078

0.00111

0.060

0.00199

0.646

0.02214

487.0

6.66

590.8

72.21

506.3

13.66

FY15-49-24

143

35

0.25

0.230

0.00323

0.082

0.00192

2.616

0.06484

1333.7

16.91

1253.7

45.53

1305.0

18.21

FY15-49-25

401

146

0.36

0.076

0.00126

0.056

0.00145

0.589

0.01851

471.5

7.56

450.0

57.40

470.3

11.83

FY15-49-26

245

55

0.23

0.071

0.00165

0.057

0.00212

0.559

0.02453

442.2

9.91

494.5

83.33

450.9

15.98

FY15-49-27

299

96

0.32

0.082

0.00120

0.057

0.00172

0.652

0.02218

509.0

7.15

501.9

66.66

509.8

13.64

FY15-49-28

515

244

0.47

0.085

0.00109

0.055

0.00113

0.650

0.01529

525.8

6.50

420.4

46.29

508.8

9.41

FY15-49-29

761

260

0.34

0.069

0.00095

0.055

0.00107

0.521

0.01199

428.9

5.71

394.5

44.44

425.7

8.00

FY15-49-30

133

53

0.40

0.352

0.00565

0.124

0.00250

6.060

0.16365

1945.4

26.91

2009.3

35.80

1984.6

23.54

FY15-49-31

701

544

0.78

0.088

0.00118

0.059

0.00106

0.718

0.01494

545.1

7.00

564.9

43.51

549.7

8.83

FY15-49-32

501

273

0.54

0.080

0.00113

0.059

0.00146

0.655

0.01836

494.4

6.75

583.4

53.69

511.8

11.26

FY15-49-33

153

134

0.87

0.138

0.00170

0.067

0.00150

1.274

0.03061

832.2

9.65

838.9

46.30

834.0

13.67

Table 6 Spot

U

Th Th/U

206

Pb/238U

1σ

207

Pb*/206Pb*

1σ

207

Pb*/U235

1σ

206

Pb/238U(Ma) 1σ

207

Pb/206Pb(Ma)

1σ

207

Pb/235U(Ma)

1σ

FY15-51-01

621

5

0.01

0.041

0.00051

0.054

0.00187

0.312

0.01023

261.4

3.15

372.3

77.77

275.8

7.92

FY15-51-02

1162

7

0.01

0.043

0.00071

0.055

0.00155

0.328

0.01021

272.2

4.38

466.7

60.18

287.9

7.80

FY15-51-03

532

6

0.01

0.044

0.00067

0.053

0.00174

0.326

0.01049

279.8

4.12

316.7

75.92

286.5

8.04

FY15-51-04

582

7

0.01

0.046

0.00076

0.055

0.00167

0.353

0.01061

289.9

4.70

416.7

68.51

307.0

7.96

FY15-51-05

1020

8

0.01

0.046

0.00078

0.053

0.00132

0.341

0.00939

290.0

4.80

344.5

55.55

298.0

7.11

FY15-51-06

806

7

0.01

0.046

0.00050

0.054

0.00223

0.345

0.01384

290.0

3.10

388.9

97.21

300.9

10.45

FY15-51-07

1011

15

0.01

0.046

0.00144

0.056

0.00558

0.337

0.01571

292.5

8.86

435.2

224.81

294.9

11.93

FY15-51-08

448

5

0.01

0.046

0.00053

0.059

0.00251

0.381

0.01512

292.5

3.24

588.9

123.13

327.5

11.12

FY15-51-09

796

5

0.01

0.047

0.00047

0.053

0.00206

0.341

0.01324

294.6

2.89

322.3

88.88

297.9

10.03

FY15-51-10

1121

9

0.01

0.047

0.00042

0.056

0.00194

0.366

0.01218

295.0

2.58

464.9

75.92

316.4

9.06

FY15-51-11

1210

12

0.01

0.048

0.00044

0.053

0.00177

0.347

0.01119

300.6

2.68

322.3

75.92

302.8

8.43

FY15-51-12

1077

9

0.01

0.048

0.00045

0.054

0.00190

0.356

0.01231

300.6

2.79

364.9

79.62

309.2

9.22

FY15-51-13

857

5

0.052

0.00075

0.053

0.00138

0.383

0.01066

327.9

4.60

324.1

59.25

329.0

7.83

FY15-51-14

939

8

0.055

0.00120

0.056

0.00143

0.428

0.01397

343.2

7.32

477.8

57.40

362.1

9.93

FY15-51-15

715

5

0.057

0.00075

0.060

0.00224

0.470

0.01432

359.5

4.60

609.3

81.47

391.5

9.89

FY15-51-16

700

5

0.01

0.048

0.00066

0.048

0.00145

0.320

0.01039

302.9

4.06

101.9

67.59

282.0

7.99

FY15-51-17

811

5

0.01

0.058

0.00129

0.062

0.00174

0.502

0.01573

364.1

7.86

733.3

60.33

412.8

10.64

FY15-51-18

510

11

0.02

0.058

0.00099

0.057

0.00166

0.455

0.01399

364.2

6.04

476.0

64.81

381.1

9.76

FY15-51-19

1024

8

0.01

0.059

0.00081

0.053

0.00157

0.437

0.01398

368.2

4.91

331.5

66.66

367.8

9.88

FY15-51-20

676

4

0.01

0.059

0.00135

0.055

0.00139

0.447

0.01432

368.8

8.21

466.7

57.40

375.2

10.05

FY15-51-21

780

24

0.03

0.060

0.00076

0.061

0.00210

0.509

0.01845

373.4

4.60

638.9

78.69

418.0

12.41

FY15-51-22

701

37

0.05

0.061

0.00144

0.061

0.00163

0.513

0.01571

381.9

8.74

655.6

57.40

420.2

10.55

FY15-51-23

524

12

0.02

0.061

0.00085

0.053

0.00157

0.449

0.01400

383.3

5.18

331.5

66.66

376.9

9.81

FY15-51-24

930

8

0.01

0.063

0.00105

0.055

0.00194

0.480

0.01750

391.4

6.36

427.8

77.77

398.1

12.01

FY15-51-25

885

85

0.10

0.064

0.00170

0.055

0.00173

0.500

0.02118

397.6

10.29

431.5

70.36

411.7

14.34

FY15-51-26

776

89

0.11

0.073

0.00151

0.059

0.00192

0.596

0.02065

454.7

9.10

564.9

67.58

474.4

13.14

FY15-51-27

567

82

0.15

0.083

0.00181

0.060

0.00213

0.702

0.03071

514.7

10.77

594.5

105.54

539.9

18.32

FY15-51-28

1908

268

0.14

0.087

0.00122

0.055

0.00145

0.660

0.01871

535.7

7.24

466.7

59.25

514.5

11.45

FY15-51-29

332

26

0.08

0.072

0.00095

0.059

0.00262

0.584

0.02627

446.1

5.71

564.9

96.28

466.7

16.85

FY15-51-30

718

137

0.19

0.130

0.00205

0.066

0.00178

1.184

0.03500

785.6

11.68

1200.0

55.55

793.2

16.28

FY15-51-31

457

225

0.49

0.076

0.00120

0.054

0.00158

0.566

0.01834

469.6

7.21

376.0

66.66

455.2

11.90

FY15-51-32

218

150

0.69

0.081

0.00134

0.058

0.00193

0.649

0.02264

500.4

7.98

538.9

72.21

507.7

13.94

FY15-51-33

489

223

0.46

0.077

0.00123

0.055

0.00165

0.587

0.01943

480.9

7.37

398.2

66.66

468.9

12.43

FY15-51-34

441

59

0.13

0.071

0.00148

0.055

0.00172

0.535

0.01930

441.6

8.94

398.2

74.99

435.2

12.77

FY15-51-35

581

154

0.27

0.074

0.00099

0.055

0.00149

0.559

0.01551

461.7

5.95

390.8

62.96

451.1

10.10

FY15-51-36

225

118

0.52

0.119

0.00255

0.063

0.00167

1.038

0.03485

727.1

14.66

694.5

57.40

723.1

17.36

FY15-51-37

429

385

0.90

0.085

0.00147

0.057

0.00146

0.667

0.01936

525.1

8.76

479.7

89.80

518.9

11.79

FY15-51-38

314

133

0.42

0.079

0.00125

0.057

0.00173

0.617

0.01900

489.0

7.46

483.4

66.66

487.8

11.93

FY15-51-39

397

144

0.36

0.078

0.00154

0.056

0.00175

0.610

0.02172

485.5

9.18

435.2

70.36

483.3

13.70

FY15-51-40

819

676

0.83

0.077

0.00112

0.058

0.00149

0.620

0.01487

478.2

6.69

516.7

57.40

490.0

9.32

FY15-51-41

837

236

0.28

0.152

0.00218

0.068

0.00144

1.428

0.03401

910.3

12.20

857.4

44.45

900.6

14.23

FY15-51-42

815

414

0.51

0.146

0.00292

0.072

0.00170

1.452

0.03822

880.8

16.41

988.9

48.15

910.8

15.83

FY15-51-43

87

23

0.26

0.214

0.00590

0.092

0.00299

2.765

0.12166

1250.8

31.30

1533.3

61.27

1346.2

32.82

Table 7 206

Pb/238U

207

Pb*/206Pb*

207

Pb*/U235

208

Pb/232Th

206

Pb/238U(Ma) 1σ

207

Pb/206Pb(Ma)

207

Pb/235U(Ma)

208

Pb/232Th

U

Th

FY15-51-01

8284

40535

0.044

0.00023

0.052

0.00104

0.314

0.00618

0.014

0.00007

276.0

1.46

287.1

46.29

277.4

4.78

271.9

1.43

FY15-51-02

5808

40878

0.045

0.00027

0.053

0.00116

0.325

0.00700

0.014

0.00009

281.5

1.66

316.7

17.59

285.6

5.37

283.7

1.73

FY15-51-03

9121

34978

0.045

0.00027

0.051

0.00107

0.319

0.00649

0.014

0.00008

283.8

1.65

253.8

48.14

280.8

5.00

278.6

1.62

FY15-51-04

9263

35772

0.045

0.00029

0.051

0.00103

0.320

0.00645

0.014

0.00009

285.0

1.81

250.1

41.66

281.7

4.97

283.9

1.75

FY15-51-05

5812

43082

0.046

0.00028

0.053

0.00124

0.332

0.00744

0.014

0.00008

287.1

1.75

324.1

21.30

291.1

5.68

287.0

1.65

FY15-51-06

6493

44967

0.046

0.00032

0.052

0.00123

0.329

0.00764

0.014

0.00008

287.3

2.00

301.9

49.07

289.2

5.84

285.2

1.65

FY15-51-07

6277

44195

0.046

0.00033

0.052

0.00158

0.328

0.01013

0.014

0.00009

288.7

2.05

279.7

68.51

288.3

7.75

290.6

1.79

FY15-51-08

5310

42551

0.046

0.00029

0.053

0.00121

0.336

0.00736

0.014

0.00008

289.2

1.78

344.5

51.85

294.1

5.60

283.0

1.51

FY15-51-09

4936

53027

0.046

0.00037

0.055

0.00122

0.348

0.00754

0.014

0.00009

289.9

2.31

413.0

50.00

303.3

5.68

285.3

1.72

FY15-51-10

6793

40938

0.046

0.00031

0.054

0.00113

0.342

0.00689

0.014

0.00009

290.1

1.92

364.9

48.14

298.5

5.22

290.9

1.73

FY15-51-11

5906

44361

0.046

0.00036

0.052

0.00116

0.329

0.00743

0.015

0.00009

290.2

2.23

279.7

51.85

289.2

5.68

293.3

1.89

FY15-51-12

5521

37842

0.046

0.00033

0.054

0.00121

0.341

0.00768

0.015

0.00009

291.1

2.05

353.8

51.85

298.3

5.82

291.2

1.70

FY15-51-13

6230

43786

0.046

0.00031

0.053

0.00101

0.337

0.00633

0.014

0.00009

291.1

1.94

324.1

44.44

294.9

4.81

288.5

1.75

FY15-51-14

5464

50715

0.046

0.00034

0.050

0.00127

0.321

0.00810

0.014

0.00008

291.9

2.10

211.2

59.25

282.4

6.23

287.7

1.51

FY15-51-15

8602

42066

0.046

0.00030

0.052

0.00103

0.333

0.00656

0.015

0.00007

292.5

1.84

279.7

44.44

291.6

5.00

292.4

1.43

FY15-51-16

7215

45416

0.047

0.00029

0.054

0.00115

0.346

0.00774

0.014

0.00008

293.6

1.78

353.8

80.55

301.4

5.85

289.3

1.59

FY15-51-17

5940

41740

0.047

0.00029

0.054

0.00137

0.346

0.00888

0.015

0.00009

294.3

1.78

366.7

57.40

302.0

6.70

292.2

1.85

FY15-51-18

5626

43620

0.047

0.00037

0.053

0.00114

0.342

0.00710

0.015

0.00010

295.9

2.30

324.1

48.14

299.0

5.37

291.9

1.99

Spot

1σ

1σ

1σ

1σ

1σ

1σ

1σ

Highlights Altai UHT granulites documented an anticlockwise P–T path involving a post–peak ITD to near-IBC process. U–Th-Pb chronological results for metamorphic zircons and monazites show two weighted mean ages of ~390 Ma and ~280 Ma. The UHT metamorphic event was likely associated with Permian reworking and the Tarim mantle plume activity.