Geochemical and geochronological studies of the Alegedayi Ophiolitic Complex and its implication for...

Gondwana Research 18 (2010) 438–454

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Geochemical and geochronological studies of the Alegedayi Ophiolitic Complexand its implication for the evolution of the Chinese Altai

Kenny Wong a, Min Sun a,⁎, Guochun Zhao a, Chao Yuan b, Wenjiao Xiao c

a Department of Earth Sciences, The University of Hong Kong, Pokfulam, Hong Kong, Chinab Key Laboratory of Isotope Geochronology and Geochemistry, Guangzhou Institute of Geochemistry, Chinese Academy of Sciences, Guangzhou 510640, Chinac State Key Laboratory of Lithospheric Evolution, Institute of Geology and Geophysics, Chinese Academy of Sciences, Beijing 100029, China

⁎ Corresponding author. Tel.: +86 852 28592194; faxE-mail address: [email protected] (M. Sun).

1342-937X/$ – see front matter © 2010 International Adoi:10.1016/j.gr.2010.01.010

a b s t r a c t
a r t i c l e i n f o
Article history:Received 29 June 2009Received in revised form 26 January 2010Accepted 31 January 2010Available online 19 February 2010

Keywords:AltaiCentral Asian Orogenic BeltAccretionGeochemistryOphiolitic complexGeochronology

The Alegedayi Ophiolitic Complex (AOC) was discovered in the northwestern part of Altai, Xinjiang, China.Strips of mafic rocks including gabbro, diabase, pillow basalt and pyroxenite, all deformed, mostlyunderwent low-grade metamorphism, are intercalated with marine-facies sedimentary strata consisting ofshale, siltstone and chert. SHRIMP zircon dating of a metagabbro sample gave an age of 439±17 Ma. In termsof whole-rock geochemistry, the AOC is composed of three distinct groups. Whereas the majority hastransitional or enriched mid-ocean ridge basalt (T-/E-MORB) affinity, varieties with oceanic island basalt(OIB) and supra-subduction zone (SSZ) affinities were also identified. Our data show that the petrogenesis ofthis mafic complex involved interaction among depleted mantle, enriched component, recycled sediments,slab-derived fluid, and metasomatized mantle wedge. Taking into account the coeval adakites and high-Mgrocks to the southwest and a slightly older arc to the northeast, we suggest that the AOC reflects an intra-arcspreading event that occurred in response to an early Silurian ridge subduction which had profound effectson the tectonic evolution of the Chinese Altai.

© 2010 International Association for Gondwana Research. Published by Elsevier B.V. All rights reserved.

1. Introduction

Extending from the Urals in the west to the Pacific in the east andfrom Siberia in the north to the Tianshan in the south, the CentralAsian Orogenic Belt (CAOB; inset of Fig. 1) is one of the largestaccretionary orogen and a most significant site of crustal growth inthe Phanerozoic (Coleman, 1989; Jahn et al., 2000; Kovalenko et al.,2004; Jahn, 2004; Helo et al., 2006). Because of the allochthonousnature of many terranes and their complicated amalgamation,tectonic history of these terranes is very controversial. A number ofmodels have been proposed to explain the accretionary history ofthe CAOB. Sengor et al. (1993) and Sengor and Natal'in (1996)considered it as a single long-lived arc and back-arc basin system,duplicated by syn-subduction stike-slip motions as the two majorblocks (Siberia and Baltica) drifted and finally collided. Otherresearchers suggested that its formation involved processes foundin modern subduction systems — collision of microcontinents andaccretion of various edifices suck as island arcs, oceanic islands,seamounts, sediments and ophiolites (Zonenshain et al., 1990;Fedorovskii et al., 1995; Buchan et al., 2002; Badarch et al., 2002;Yakubchuk, 2004; Dobretsov and Buslov, 2004; Buslov et al., 2004a;Windley et al., 2007; Kroner et al., 2008). Many researchers agree

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that the best analogy to the evolution of the CAOB is thesubduction–accretion systems in the circum-Pacific region (Xiaoet al., 2004; Windley et al., 2007; Safonova et al., 2009; Xiao et al.,2010-this issue).

This paper reports geochemical data and zircon dating results forthe newly mapped Alegedayi Ophiolitic Complex (AOC). We suggestthat formation of the AOC and some other coeval, geochemically-special rocks (adakite, boninite and picrite) found along the southernmargin of the Chinese Altai (present coordinates) can be explained byridge subduction which is a process found in Southern Chile section ofthe Pacific subduction system (Bourgois et al., 1996; Lagabrielle et al.,2000). It is possible that our samples with both supra-subductionzone (SSZ) character and N-MORB signatures represent an intra-arcocean spreading event responding to a change in convection patternin the mantle wedge caused by a raised geotherm. Mafic rocks withenriched and transitional (E- and T-) MORB features may implymixing of depleted, sub-lithospheric mantle of the subducting platewith an enriched source with oceanic island basalt (OIB) affinities. Infact, this early Silurian ridge subduction model explains not only thenearly 20 Ma of magmatic hiatus (Tong et al., 2007) but also thesignificant change in tectonic setting and magma source of theChinese Altai (Yuan et al., 2007; Long et al., 2007; Sun et al., 2008; Xiaoet al., 2009; Sun et al., 2009). Overall, AOC serves as a gooddemonstration of the notion that SSZ ophiolite in orogenic beltsignifies production of oceanic crust during basin closure prior tomajor collision (Dilek, 2003a).

Published by Elsevier B.V. All rights reserved.

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Fig. 1. Simplified geological map of the Chinese Altai and adjacent areas (1— The Altaishan Terrane; 2— The northwest Altaishan Terrane; 3— The central Altaishan Terrane; 4— TheQiongkuer–Abagong Terrane; 5 — The Erqis Terrane; 6 — The Shaerbulake Terrane). Inlet: The Central Asian Orogenic Belt (CAOB) and geological units nearby (the star marks thelocation of the Chinese Altai).

439K. Wong et al. / Gondwana Research 18 (2010) 438–454

2. Geological background

2.1. Regional geology

The 2500 km-long NW–SE trending Altai orogen stretches fromthe Gorny Altai in Russia and the Rudny Altai in East Kazakhstan to theGobi Altai in southwestern Mongolia (Xiao et al., 1992; Badarch et al.,2002; Buslov et al., 2004b; Xiao and Kusky, 2009 and papers therein;Ao et al., 2010-this issue). It was regarded as a subduction-accretedcontinental margin of the Siberia craton formed in the Paleozoic(Coleman, 1989; Sengor et al., 1993). Though Sm–Nd isotopic dataand zircon xenocrysts from granitoids may hint at the presence of aPrecambrian basement (Hu et al., 2000, 2006; Windley et al., 2002;Wang et al., 2009), the Chinese Altai is largely made up of a long-livedPaleozoic magmatic arc with minimal ancient crustal reworking, asrevealed by recent studies on Hf and U–Pb systems of zircons frommetamorphosed granitic and sedimentary rocks (Long et al., 2007;Sun et al., 2008). The entire Chinese Altai is truncated by the Irtyshfault zone (also called the Erqis faults; the Erqishi shear zone or theErtix faults) which likely accommodated large-scale strike-slipmotion (Laurent-Charvet and Charvet, 2003) and top-to-SW thrusting(Briggs et al., 2007) during the early Permian. This large fault zoneseparates the Junggar low-relief hummocky terranes with lowergreenschist-facies rocks in the southwest from the uplifted Altai

mountain ranges with high-grade metamorphic rocks in thenortheast.

Internally, the Chinese Altai can be divided into several fault-bounded terranes (Windley et al., 2002; Fig. 1). The Altaishan Terrane(Terrane 1) is composed of Late Devonian–Early Carboniferous clasticsediments, limestones and some minor island-arc volcanic rocksmetamorphosed at lower greenschist facies (Zhuang, 1993). Grani-toids with allegeable Silurian to Early Devonian zircon U–Pb ageswere reported (Lou, 1997).

The northwest Altaishan Terrane (Terrane 2) is made up of aMiddle Ordovician turbidite sequence of lower greenschist facies(Long et al., 2007) which is unconformably overlain by LateOrdovician volcanoclastic sediments (He et al., 1990). Recent zirconU–Pb dating on Kanasi granitic pluton (Tong et al., 2007) supports theoccurrence of early Devonian igneous event.

The central Altaishan Terrane (Terrane 3) is the largest terrane andis composed predominantly of Early Paleozoic sediments metamor-phosed at medium to high grade (Sun et al., 2008). Arc-relatedmagmatic rocks of various ages (Cambrian to Devonian) andmetamorphism from greeschist- to granulite-facies are also found(Windley et al., 2002; Wang et al., 2006). Presence of Precambrianbasement is inferred from xenocrystic zircon ages of a rhyodacite anda tonalitic gneiss (Windley et al., 2002) and whole-rock Nd crustalresidence ages of amphiolites (Hu et al., 2000). It likely represents a

440 K. Wong et al. / Gondwana Research 18 (2010) 438–454

“magmatic arc built on a continental margin dominated by Neopro-terozoic rocks” (Sun et al., 2008).

The Qiongkuer–Abagong Terrane (Terrane 4), of which theAlegedayi Ophiolitic Complex locates, consists of volcaniclastic rocksof Late Silurian–Early Devonian age overlain by Mid Devonianturbiditic sandstone, pillow basalts and some siliceous volcanicrocks. Volcanic rocks and granites from this terrane are largelyDevonian but Carboniferous and Permian ones are not uncommon(Zhang et al., 2000; Briggs et al., 2007; Yuan et al., 2007). AlthoughNeoproterozoic xenocryst ages and subchondritic epsilon hafniumvalues from zircons of Cambrian–Ordovician granitoids indicatecrustal recycling, +ve εHf of zircons from Devonian granitic plutonssuggests dominance of juvenile mantle input (Sun et al., 2008).Permian igneous activity is very minimal and the 276 Ma I-A typeLamazhao pluton (Wang et al., 2005) seems to be the only example.

Bounded to the south by the Irtysh fault, the Erqis Terrane (Terrane5) contains Devonian–Late Carboniferous volcaniclastic rocks meta-morphosed at greenschist- to amphibolite-facies and large amount ofhigh-grade gneisses. The later are Ordovician–Devonian arc-relatedintrusions deformed in the Permian thrusting (Briggs et al., 2007) andcontaminated by Paleoproterozoic materials (Qu and Chong, 1991).

Across the Irtysh fault is the Shaerbulake Terrane (Terrane 6)which is composed of weakly metamorphosed (up to lower greens-chist facies) volcanoclastic rocks. Though Devonian–Carboniferousfossils (Xiao et al., 1992) were found in these rocks, latest dating ofadakites indicates Silurian igneous activity (Zhang et al., 2008). Ourunpublished result of zircon U–Pb dating on a basalt sample previousthought as a “Carboniferous unit” give Permian age and inheritance ofOrdovician and Neoproterozoic. The new data suggest that Shaerbu-lake Terrane belongs to Chinese Altai “active margin built on an oldmicro-continent” but not the young intra-oceanic subduction systemsof Eastern Junggar. In addition, presence of high-Mg rocks such asboninite and picrite suggests that this terrane is probably a forearc(Zhang et al., 2003, 2005b). On the southern margin of this terrane isthe Aermantai ophiolite with arc signature (Wang et al., 2003). It isbelieved to be an island-arc formed in the late Cambrian and lateraccreted to the Chinese Altai in Permian (Xiao et al., 2006, 2009). Thisophiolitic belt probably marks the northern margin of the easternJunggar arc-accretionary system.

Fig. 2. a. Aerial photo of the field area (image captured from G

2.2. Field geology

Along the faulted zone of Terranes 4, we discovered a NW–SEtrending ophiolitic belt near the Alegedayi river. Viewed from the sky(Fig. 2a), this 35 km-long, 2.5 to 5 km-wide Alegedayi OphioliticComplex (AOC) exists as a dark band standing out from the pinkgranites and the yellowishwhite desert. As shown in Fig. 2b, the belt issandwiched by early Devonian granitoids — Tarlang Batholith to theNE (Yuan et al., 2007) and the Hahahe Granodiorite to the SW (Yuanet al., 2006). The contact between the belt and the Tarlang Batholith isa shear zone which is filled by many quartz veins and thus appears asa white band (Fig. 3a). The AOC was affected by both top-to-SWthrusting (Briggs et al., 2007) and sinistral shearing (Laurent-Charvetand Charvet, 2003) in the Early Permian.

Field mapping reveals that the AOC is composed of alternatingstrips of mafic rocks and marine-facies sedimentary rocks dominatedby shale, siltstone and some red chert beds (Fig. 3b). While gabbrosand pyroxenites can be found, dolerites and basalts aremost common.Despite the deformation, pillow structure is still preserved at somelocalities (Fig. 3c). Serpentinization is restricted to some outcropsonly, and most of the rocks experienced only low-grade metamor-phism as exemplified by greenschists and phyllites. A few thin layersof plagioclase-rich volcaniclastic rocks were also found intercalated inthe strata, inferring a high proximity to a felsic volcanic source.

Though the AOC is tectonically disturbed, several lines of evidencesuggest that it is not an ophiolitic mélange. Our structural analysesshow that the repetitive layers of mafic and sedimentary rocks have avery uniform sub-vertical dips belonging to an upright, tight foldsystem (Fig. 3a) which was formed in one single deformational event.Moreover, unlike Franciscan-type ophiolites (Dilek, 2003b), there isneither block of non-ophiolitic lithology nor fragment of exoticgeological units. Also, the AOC is quite a distance from the proposedsuture between the Chinese Altai active continental margin and theEast Junggar island-arc system. The presence of Proterozoic materialand early Paleozoic arc-related rocks to the SW (Terrane 5) and NE(Terrane 3) of the AOC supports the idea that this part of the ChineseAltai is an intra-arc setting (Briggs et al., 2007). Overall, the AOC is notan ophiolitic mélange composed of rocks of different origins but adismembered ophiolite formed in an intra-arc environment.

oogle Map). b. Simplified geological map of the field area.

Fig. 3. a. Top-to-SW thrust contact between the granitoid and the tightly folded Alegedayi Ophiolitic Complex (the view is due South). b. Outcrop of the sedimentary rock layersshowing two phases of deformation. c. Pillow structure preserved in the folded basalt layers.


3. Sample description and analytical techniques

Mafic rocks analyzed in this study include basalts, dolerites andgabbros which were metamorphosed up to greenschist facies.Although some igneous textures and structures can still be observedin the field and under the microscope, the mineral assemblages ofthese samples are mainly metamorphic−amphibole+sodic plagio-clase+biotite+chlorite+quartz+/−epidote. To leach out alter-ation minerals, the samples are crushed into small chips and soakedin cold 6 N hydrochloric acid before grinded into powders.

Major and trace element analyses were performed in theDepartment of Earth Sciences, The University of Hong Kong.BHVO-2 and AGV2 were used as reference materials, and data are

presented in Table 1. Major oxide compositions were determinedby X-ray fluorescence (XRF) spectroscopy on fused glass beadsusing a Philips PW 2400 spectrometer. We followed the proceduresdescribed by Norrish and Hutton (1969) for matrix correction. Theanalytical errors are about 1–2% for SiO2, Al2O3, TiO2, Fe2O3 andMgO and less than 3% for the other oxides. Trace elementsconcentrations were determined using VG Plasma-Quad Excellinductively coupled plasma mass spectrometry (ICP-MS). 103Rhwas added to the sample solutions as an internal standard tocorrect the effect of drifting. Procedures for sample preparation,powder digestion and external calibration using multi-elementstandard solutions are described by Qi et al. (2000). The accuraciesfor most elements are better than 5%.

Table1

Major

andtraceelem

entco

mpo

sition

sof

theAlege

dayi

Oph

olitic

Complex

.

Sample

MORB

-typ

eOIB-typ

eSS

Z-type

Subg

roup

ASu

bgroup

B

OP4

2HB0

1HB0

4HB0

5HB0

6HB1

0HB1

1OP1

9OP2

0OP2

1OP2

4OP2

6OP3

9BE

J149

HB2

2HB2

3BE

J148

BEJ151

BEJ152

BEJ154

BEJ155

BEJ156

BEJ157

BEJ158

BEJ159

BEJ160

SiO2

47.20

47.39

47.33

48.28

45.98

49.96

47.97

44.35

46.00

46.37

47.41

46.68

48.99

50.36

47.35

47.09

45.82

47.36

45.76

50.37

50.52

50.14

47.06

48.38

47.05

48.06

Al 2O3

15.61

16.93

16.20

16.12

15.54

15.91

15.80

14.93

15.84

15.60

14.63

15.67

13.91

15.67

16.22

16.03

18.19

16.12

16.90

17.67

16.63

17.10

18.13

17.32

17.97

17.14

CaO

12.22

9.65

10.06

9.36

10.32

9.83

10.35

12.14

13.58

14.00

12.90

15.42

11.55

7.86

7.24

7.69

10.99

10.63

10.56

7.11

8.60

8.48

11.18

11.86

10.84

11.58

Fe2O3

10.28

10.77

11.31

11.16

12.31

9.55

10.90

13.68

12.10

11.56

12.31

10.74

11.63

11.19

12.96

13.43

10.49

12.86

12.83

10.54

10.40

10.31

9.61

9.11

9.78

9.10

K2O

0.24

0.18

0.14

0.13

0.13

0.12

0.11

0.20

0.16

0.13

0.28

0.10

0.45

0.43

0.16

0.17

0.24

0.17

0.13

0.07

0.07

0.09

0.37

0.30

0.26

0.17

Na 2O

1.83

3.48

2.75

3.00

2.38

3.19

2.57

2.02

2.39

1.41

3.00

2.10

2.73

3.48

3.88

3.86

2.41

2.52

2.54

4.88

4.30

4.54

2.21

2.44

2.69

3.48

MgO

9.93

6.90

7.83

7.49

8.53

7.28

8.16

9.24

6.91

7.62

6.71

5.03

7.32

6.70

6.27

6.11

9.61

8.01

9.01

7.51

7.62

7.62

8.90

8.50

9.13

8.40

MnO

0.14

0.14

0.14

0.15

0.15

0.12

0.14

0.17

0.13

0.17

0.18

0.13

0.18

0.13

0.18

0.18

0.15

0.21

0.19

0.15

0.16

0.15

0.14

0.15

0.15

0.16

FeOT

9.25

9.69

10.17

10.04

11.08

8.59

9.81

12.31

10.89

10.40

11.07

9.66

10.46

10.07

11.66

12.08

9.44

11.57

11.55

9.49

9.36

9.28

8.65

8.20

8.80

8.18

P 2O5

0.10

0.25

0.24

0.24

0.25

0.21

0.20

0.21

0.27

0.14

0.21

0.25

0.15

0.47

0.55

0.52

0.17

0.16

0.17

0.15

0.15

0.14

0.16

0.14

0.17

0.18

TiO2

1.07

1.99

1.96

2.00

2.01

1.82

1.79

1.81

1.59

1.60

1.42

1.50

1.42

2.33

3.13

3.16

1.55

1.60

1.66

1.35

1.32

1.23

1.50

1.38

1.46

1.41

LOI

1.37

2.33

2.04

2.07

2.39

2.01

2.00

1.25

1.04

1.40

0.94

2.39

1.67

1.37

2.09

1.77

0.38

0.35

0.24

0.20

0.23

0.21

0.73

0.43

0.50

0.30

Mg#

65.68

55.92

57.84

57.06

57.84

60.17

59.72

57.23

53.07

56.65

51.91

48.15

55.49

54.26

48.94

47.40

64.48

55.25

58.16

58.51

59.19

59.41

64.72

64.88

64.92

64.67

Ti63

8711

939

1174

712

005

1206

110

913

1073

710

865

9511

9573

8524

8974

8537

1598

218

737

1893

693

0396

1799

4180

8579

2373

7489

8882

8487

5484

68Ni

218.6

92.4

113.3

93.4

114.1

125.5

121.7

299.3

125.7

172.8

45.3

75.4

62.6

55.8

65.3

64.7

192.2

189.8

191.7

139.7

133.4

116.4

204.7

184.1

182.3

171.4

Y24

.79

38.34

35.77

36.96

35.69

37.62

32.36

36.49

38.15

34.73

34.88

37.10

25.74

38.84

39.12

40.21

22.13

28.64

31.02

32.94

30.69

31.97

21.74

20.48

22.66

23.55

Zr84

.67

121.90

104.70

118.80

102.45

102.75

98.75

107.50

107.80

71.50

71.30

89.45

67.25

155.68

202.55

197.10

104.68

104.60

87.10

119.95

105.80

105.45

114.84

98.01

113.88

113.02

Nb

1.38

7.89

7.45

7.92

7.32

4.57

3.79

5.52

7.33

5.94

3.23

7.06

4.05

21.02

50.05

53.25

1.58

2.35

2.24

2.90

2.63

2.57

1.63

1.54

1.81

1.56

La2.08

8.03

7.90

8.10

7.69

6.75

5.71

8.09

10.43

9.31

8.36

11.22

6.60

16.05

27.39

27.75

3.60

4.50

3.36

6.41

5.48

5.63

3.27

3.76

3.74

3.97

Ce4.47

19.85

18.63

18.92

18.41

17.60

14.50

20.07

26.05

22.00

18.97

26.26

14.70

35.95

55.55

57.15

13.43

13.04

11.28

15.91

15.66

13.67

13.92

11.49

13.04

13.35

Pr1.18

3.21

3.01

3.11

3.01

2.95

2.51

3.35

3.94

3.41

3.04

3.90

2.50

5.42

7.14

7.20

2.17

1.82

1.62

2.03

2.18

2.05

1.83

1.88

2.07

2.00

Nd

7.31

17.17

16.32

16.71

16.32

16.24

14.26

18.90

20.86

18.39

16.55

21.97

13.86

22.37

32.38

32.80

11.79

10.97

9.85

10.95

11.71

11.43

9.61

10.37

11.08

10.93

Sm2.65

4.54

4.16

4.37

4.38

4.38

3.87

5.31

5.53

4.76

4.75

5.58

4.02

6.10

6.90

6.67

3.48

3.32

3.77

3.58

3.20

3.54

3.10

2.82

3.47

3.19

Eu1.11

1.72

1.45

1.47

1.45

1.54

1.43

1.91

2.24

1.78

1.67

2.67

1.42

2.08

2.16

2.21

1.37

1.36

1.46

1.37

1.03

1.37

1.17

1.20

1.53

1.32

Gd

3.62

4.85

4.58

4.76

4.64

4.86

4.45

6.15

6.25

5.65

5.46

6.27

4.52

6.33

6.80

6.36

3.39

4.04

3.59

4.10

3.57

4.60

3.40

2.98

4.00

3.55

Tb0.65

0.82

0.72

0.78

0.78

0.80

0.74

1.06

1.04

0.98

0.93

1.01

0.77

1.03

0.98

0.95

0.63

0.79

0.68

0.75

0.67

0.80

0.67

0.58

0.65

0.66

Dy

4.94

5.52

5.04

5.31

5.19

5.49

5.11

7.17

7.20

6.96

6.66

7.03

5.17

6.62

6.06

5.95

4.21

4.63

4.49

5.10

4.19

4.98

4.34

3.88

4.63

3.89

Ho

1.05

1.15

1.06

1.06

1.09

1.14

1.08

1.49

1.52

1.42

1.42

1.46

1.08

1.37

1.21

1.21

0.88

1.01

1.02

1.03

0.92

0.95

0.84

0.90

1.02

0.85

Er2.91

3.26

2.99

3.14

3.12

3.11

2.97

4.23

4.32

3.97

3.93

4.22

3.07

3.49

3.30

3.27

2.37

2.48

2.50

2.79

2.78

2.69

2.34

2.30

2.83

2.39

Tm0.43

0.46

0.41

0.46

0.43

0.46

0.44

0.66

0.65

0.60

0.60

0.61

0.45

0.48

0.45

0.44

0.36

0.34

0.36

0.42

0.42

0.35

0.33

0.33

0.40

0.35

Yb2.91

2.86

2.60

2.81

2.73

2.78

2.80

4.15

4.35

3.77

3.94

3.81

3.03

3.14

2.82

2.75

2.12

2.28

2.38

2.53

2.84

2.22

1.98

1.97

2.19

2.20

Lu0.42

0.41

0.36

0.39

0.37

0.37

0.37

0.58

0.56

0.56

0.56

0.53

0.41

0.48

0.40

0.37

0.35

0.38

0.34

0.38

0.40

0.36

0.36

0.31

0.33

0.35

Hf

1.75

2.57

2.07

2.34

2.15

2.15

2.23

3.04

2.82

2.42

2.18

2.54

2.01

3.93

4.30

3.78

2.25

2.68

2.15

2.86

2.93

2.54

2.11

2.33

2.28

2.48

Ta0.11

0.50

0.43

0.45

0.43

0.30

0.28

0.35

0.43

0.39

0.21

0.42

0.25

1.35

2.67

2.68

0.12

0.16

0.13

0.22

0.18

0.15

0.13

0.12

0.13

0.12

Th0.11

0.50

0.44

0.44

0.44

0.40

0.32

0.47

0.65

0.49

0.81

0.66

0.64

1.07

1.86

1.75

0.13

0.20

0.13

1.78

1.24

1.43

0.13

0.15

0.14

0.13

U0.05

0.17

0.15

0.16

0.16

0.12

0.13

0.20

0.26

0.57

0.43

0.55

0.20

0.32

0.63

0.61

0.08

0.11

0.15

0.66

0.62

0.55

0.06

0.09

0.07

0.08

ΔNb

−0.48

0.09

0.16

0.10

0.17

−0.01

−0.12

0.02

0.16

0.37

0.11

0.29

0.14

0.32

0.48

0.54

−0.70

−0.43

−0.26

−0.39

−0.36

−0.35

−0.77

−0.69

−0.70

−0.75


Fig. 4. Concordia plot of SHRIMP U–Pb zircon analytical results for the metagabbro(OP19) of the Alegedayi Ophiolitic Complex.


A metagabbro (OP19) was chosen for zircon dating. Throughstandard heavy liquid and magnetic techniques, zircons wereseparated and non-magnetic grains with a size N25 μm were hand-picked and fixed on an epoxy resin mount. The surface of the mountwas then polished until most of the zircons had their centers exposed.While their morphology and distribution were recorded by photostaken under reflected- and transmitted-light, their U–Th–Pb isotopeswere analyzed using the WA Consortium Sensitive High ResolutionIonMicroprobe (SHRIMP) II at Curtin University of Technology (Perth,Australia) with 6-scan duty cycles. The Sri Lankan gem zircon (CZ3)was used as a standard to monitor isotopic ratios. The decay constantsrecommended by Steiger and Jager (1977) and formula summarizedby Faure and Mensing (2005) were used to calculate the ages of oursamples. Data reduction was completed with the use of softwareSQUID and ISOPLOT developed by Ludwig (2001a,b) respectively.Individual data presented in this study are shown with 1σ errorwhereas the weighted mean ages are quoted at 95% confidence level(2σ).

4. Age of the AOC

As metagabbro is the most abundant rock type in the AlegedayiOphilitic Complex, a sample (OP19) from this group was chosen forgeochronological study, and the results are presented in Table 2.SHRIMP analyses of zircons yielded concordant or almost concordant206Pb/238U ages ranging from 408±25 Ma to 455±13 Ma with aweighted mean of 439±17 Ma (MSWD=2.0) (Fig. 4). As theirelongated prismatic morphologies, small sizes and relatively highTh/U ratios (0.41–1.42) suggest an igneous origin for these zircons(Williams and Claesson, 1987; Hoskin and Black, 2000; Hartmannand Santos, 2004), we interpret that the AOC was formed in theearly Silurian.

5. Geochemistry and petrogenesis of the AOC

5.1. Effects of alteration

Despite the fact that only samples without significant visibleweathering were selected for chemical analyses, a wide range of loss-on-ignition (LOI; 0.18–4.63 wt.%) indicates that some samples weresubject to considerable alteration by either sea-water or metamorphicfluid. Though samples with LOI higher than 2.5 wt.% were excluded soas to avoid effects of alteration, we are very cautious in using mobileelements like uranium (U) and large ion lithophile elements (LILE)such as barium (Ba), rubidium (Rb), strontium (Sr) and potassium (K)for petrogenetic interpretation.

Apart from thorium (Th) which is considered to be relativelyimmobile during low-grade metamorphism (Roser and Nathan,1997), other high field strength elements (HFSE) demonstrateconsistence within a particular type of rock, suggesting that chemicalcompositions of our samples were not seriously modified by post-magmatic processes. Rare earth elements (REE) are relativelyimmobile and unfractionated during weathering, hydrothermalalteration and low-grade metamorphism (Rollision, 1993; Daux

Table 2Abundance, isotope ratios and ages of the sample OP19.

Spot U(ppm)

Th(ppm)

232Th/238U

206Pb*(ppm)

204Pb/206Pb

%206Pbc 207Pb*/206Pb*

OP19-11 697 959 1.42 45.91 0.00816 14.7 0.082OP19-5 464 182 0.41 27.45 0.00161 2.9 0.048OP19-8 149 102 0.71 9.22 0.00094 1.7 0.061OP19-7 287 158 0.57 18.59 0.00310 5.6 0.069OP19-9 379 303 0.83 24.01 0.00066 1.2 0.062OP19-1 163 148 0.94 10.37 0.00056 1.0 0.054

Remarks: Pb* = radiogenic lead; Pbc = common lead; err = error.

et al., 1994). The sub-parallel REE patterns, hence, likely representoriginal trends. In general, our observations indicate that HFSE andREE in our mafic rocks were resistant to sea-water alteration (Luddenand Thompson, 1978; Bienvenu et al., 1990). The overall coherence ofour samples implies that their geochemical characteristics wereinherited from their precursors.

5.2. Major and trace element compositions

All analyzed samples are basaltic in composition with silica contentsranging from 44 to 51 wt.% (Fig. 5a; Le Maitre et al., 1989). They are allolivine-normative, and most have, in addition, nepheline in the CIPWnorm. Two truly alkali rocks are distinguished by the use of Nb/Y−Zr/(P2O5⁎10,000) diagram (Fig. 5b; Floyd and Winchester, 1975). TheFeOT/MgO vs SiO2 diagram (Fig. 5c; after Miyashiro, 1974) reveals thatall sub-alkaline samples are tholeiitic rather than calc-alkaline.Magnesium contents of the basalts are moderate to high (∼5–12 wt.%)and Mg-numbers [Mg#=100⁎Mg/(Mg+Fe) molar ratio] range from47 to 67. The two alkali basalts exhibit significantly higher TiO2 contents(∼3.1 wt.%)when comparedwith the rest of our samples (1 to 2.3 wt.%).Plots of major elements against MgO (Fig. 6a–e) do not show a cleartrend for all samples, implying that fractional crystallization of particularphases such as ferromagnesian minerals, Ti-bearing opaques andplagioclase were not major factors governing petrogenetic relations ofour samples. However, removal of olivine is a common process asindicated by the positive correlation between MgO contents andconcentrations of nickel (Ni) (Fig. 6f).

As what the ternary plot of Zr–Th–Ta (Fig. 7; Wood, 1980) shows,most of our samples fall in the field of N-MORB and E-MORB whereasthe rest scatter in two other fields: OIB and volcanic-arc basalt.Further examination reveals that the studied samples can be dividedinto three groups: 1) mid-ocean ridge basalt (MORB)-type, 2) oceanisland basalt (OIB)-type and 3) supra-subduction zone (SSZ)-typewhich has two subgroups.

% err 207Pb*/235U

% err 206Pb*/238U

% err errcorr

206Pb/238U Age

1σerr

82.6 0.740 82.9 0.065 6.4 0.077 408.3 25.212.5 0.447 12.7 0.067 2.5 0.193 417.6 9.96.4 0.591 6.8 0.071 2.5 0.367 439.5 10.7

12.9 0.677 13.2 0.071 2.8 0.215 443.3 12.17.3 0.625 7.7 0.073 2.4 0.313 453.5 10.65.7 0.548 6.4 0.073 2.9 0.449 455.4 12.6

Fig. 5. a) Silica vs. total Alkalis (TAS) diagram (Le Maitre et al., 1989) of all samples;b) Nb/Y−Zr/(P2O5⁎10,000) diagram (after Floyd andWinchester, 1975) and c) SiO2 vsFeOT/MgO of all sub-alkaline samples (after Miyashiro, 1974).


The OIB-type samples distinguish themselves from the other twogroups by their significant enrichment in incompatible elements(Fig. 8a). Their trace element variation patterns are similar to that ofaverage OIB (Sun and McDonough, 1989) and they show higher HFSEand LILE contents than average enriched MORB (E-MORB, Sun andMcDonough, 1989). Their most distinctive feature — a positive Nb–Taanomaly— is reflected by their high Nb/La values (1.83 to 1.92) whichstand out among all samples and are even higher than average OIB(1.50; Saunders and Tarney, 1984). Their high Th/Sm ratios (∼0.265)suggest that they are enriched in LILE. Also, these OIB-type samplesare LREE enriched ((La/Sm)N∼2.6 and (La/Yb)N∼7), resembling thatof average OIB (Sun and McDonough, 1989). They have REE contents10 times (for HREE) to 100 times (for LREE) higher than those of C1chondrite (Fig. 9a). Their heavy REEs (HREE) are depleted withrespect to MORB. As their low Mg# (average 48) and Ni contents (65)

suggest, these alkaline lavas experienced fractionation of olivine andthus do not represent primary melts.

The MORB-type group displays a compositional spectrum rangingfrom N-MORB through T-MORB to E-MORB. The major differencebetween these two varieties is demonstrated by their REE patterns(Fig. 9b). As for their LREE sections, the N-MORB samples havepositive slopes (0.506b(La/Sm)Nb1), whereas the T- and E-MORBdisplay sub-horizontal to slightly negative slopes (1.02b(La/Sm)Nb1.29). The steepness of HREE sections, however, is more or less thesame (1.31b(Sm/Yb)Nb2.21). In general, our MORB samples haveenriched trace element contents relative to the average N-MORBvalues (Sun and McDonough, 1989; Fig. 8b). Excluding mobileelements (Rb, Ba, K and Sr) with highly variable abundances, thedegrees of enrichment in incompatible elements are confinedbetween 1 (Ti) time and 10 (Th) times of the average N-MORB andare noticeably weaker than those of the OIB-type samples. Troughs inHFSE such as Nb, Ta, Zr, Hf and Ti can be found in some samples,possibly indicating a subduction influence. MORB-type samplesgenerally have Nb/La value close to that of average MORB (0.83;Saunders and Tarney, 1984). All MORB-type samples have Th/Smhigher than average N-MORB (0.046) but lower than average E-MORB(0.23) (Sun and McDonough, 1989).

The remaining samples belong to the SSZ-type group. Thesesamples show characteristic depletions in Nb and Ta relative to REE(Thirlwall et al., 1994) and also progressive enrichment in traceelements with increasing incompatibility (Fig. 8c). The most prom-inent feature of these rocks, namely a negative Nb–Ta anomaly, isquantified by the Nb/La values (0.394–0.668) which are the lowestamong all groups. In fact, these rocks can be further divided into twosubgroups. Subgroup A samples have the highest Th/Sm among allsamples (0.38–0.49), imitating the elevated LILE/LREE ratio of arcmagma (Tatsumi et al., 1986). As shown in Fig. 9c, their REEconcentrations are about 10 to 30 times of those of C1 chondrite.The slightly negative slopes of their REE patterns indicate relativeenrichment of LREE over HREE (1.38b(La/Yb)Nb1.82) which isanother diagnostic feature of arc magmas. The extent of such en-richment is higher than that of T-MORB samples (1.01b(La/Yb)Nb1.77) but lower than those of E-MORB and OIB-type populations(2.02b(La/Yb)Nb7.23). With respect to average N-MORB they are, toa very small degree, depleted in HREE (N0.7 times) and enriched inLREE (b3 times). Subgroup B samples, on the other hand, have flat REEpatterns similar to those of basalts from marginal basins like Mariana(Gribble et al., 1996, 1998) and Okinawan (Shinjo et al., 1999), andthey show no LILE enrichment (Th/Sm∼0.05).

5.3. Source characteristics and melting processes

5.3.1. Mantle source and partial meltingThe first question to address here is whether the varied magma

composition, ranging from very depleted to highly enriched, is theeffect of source enrichment or the result of various degrees of melting.Fitton et al. (1997, 2003) proposed a criterion, the ΔNb (whereΔNb=1.74+log(Nb/Y)−1.92log(Zr/Y)), to define mantle enrich-ment and successfully demonstrated that its significance is of globalscale. Since ΔNb is a fundamental source characteristic and isinsensitive to magma processes such as fractional crystallization,this parameter is used for identification of an enriched mantle source.As shown in Fig. 10a (after Condie, 2005), the OIB-type samples, E-MORB and many T-MORB have +ΔNb (Table 1), suggesting anenriched mantle origin. Furthermore, these T-MORB samples showtwo trends. The more obvious one parallels the lower boundary of theIceland array (ΔNb=0) and represents differences in the degree ofpartial melting and/or source depletion through melt extraction.Conversely, other T-MORB samples plot on the boundary or evenwithin the field of N-MORB, forming a trend of decreasing ΔNb valuesfrom positive to negative. Given the intrinsic nature of ΔNb, such

Fig. 6. a–f) MgO plotted against SiO2, TiO2, FeOT, MnO, Al2O3/CaO and Ni respectively.


transition can only be explained by mixing between the enrichedmantle and a−ΔNb source, i.e. the depletedmantle. Presence of a fewT-MORB and SSZ-type samples in the overlapping area of the Arc andNMORB fields hints that sea-floor spreading happened over asubduction zone.

Another approach to investigate the nature of mantle source(s) isthrough melt modeling. As shown in the La/Yb vs. Zr/Nb diagram(Fig. 10b; after Aldanmaz et al., 2008), more than half of the T-MORBsamples locate at b0.1% of depleted mantle (DM). Since such a degreeof partial melting is unrealistically low, our rocks could not beproduced by a single-stage melting of depleted mantle. Anothercandidate is primitivemantle, but this faces twomajor problems. First,Zr/Nb of these rocks are higher than that of PM itself; though one mayargue that their high ratios may be result of contamination with arc-related components, as inferred from their subtle drifting towardsSSZ-type samples. Second, even if we exclude samples with possibleinfluence of subduction-derived material (those outside the MORBfield), the degree of partial melting modeled would be unreasonably

Fig. 7. Zr–Th–Ta plot (after Wood, 1980).

large (40–50%). In fact, if PM was the source, large differences in LREEconcentrations among our samples would only be attributed tovariation in degree of batch melting from 1 to 10% (fig. 4.36 ofRollision, 1993) which is much lower than the model suggests.Though we cannot totally rule out the involvement of PM, mixingbetween a more enriched component and a depleted source isrequired to explain the wide range of Zr/Nb and La/Yb of the T-MORBsamples. On the other hand, the OIB-type samples are totally out ofreach of the model curve of primitive mantle, so an even moreenriched source is required.

More insights intomelting processes of our samples come from theNb/Zr vs. Zr/Y diagram (Fig. 10c; after Ichiyama et al., 2008). Given asimilar degree of partial melting (1–2%), our OIB samples showsignificantly lower Zr/Y than those of alkali basalts from St. Helena.Similarly, our E-MORB samples are quite different from the tholeiitesof Hawaii. It can thus be concluded that enriched components of theAlegedayi ophiolite are not chemically comparable to magmas fromocean islands (sensu stricto) which are always related to plumes. Asthe model curves delineate, our OIB-type samples may originate froma relatively shallow mantle source with spinel as a more dominantphase over garnet. Comparatively, T-MORB samples are morecomplicated in petrogenesis. These samples resemble tholeiitesfrom three different settings — Iceland, Ontong Java and East PacificRise. The scatter of these data points across several fields substantiatesthe mixing model. Furthermore, their distribution around 20% of thetwo lherzolite model curves hints at a relatively higher degree ofpartial melting and thus a low pressure environment with a thinoverlying lithosphere, i.e. young oceanic crust or spreading center.

Overall, our trace element results suggest that the AOC cannot beproduced by melting of depleted or primitive mantle alone. As themain body of the AOC, the T-MORB samples were possibly formed bymixing of melts derived from both depleted (−ΔNb) and enriched(+ΔNb) sources represented by the N-MORB and OIB-type samplesof this study. The blending was obviously not complete since theseMORBs show high variability in the degree of enrichment. Addition-ally, the existence of depleted end-members in the MORB spectrumand the predominance of tholeiite indicate that there was consider-able contribution from asthenospheric mantle melted at shallow

Fig. 8. N-MORB normalized multi-element patterns for the a) OIB, b) MORB and c) SSZ.(E-MORB, OIB compositions and N-MORB normalization values are from Sun andMcDonough, 1989).

Fig. 9. Chondrite normalized REE patterns for the a) OIB, b) MORB and c) SSZ. (N-MORB,E-MORB, OIB compositions and C1 chondrite normalization values are from Sun andMcDonough, 1989).


depth. This lherzolite source probably underwent high-degreemelting in between spinel and garnet stability fields when itencountered the alkaline melt with OIB character. Consideredcollectively, these observations suggest that the mafic rocks did notoriginate from a typical oceanic island setting inwhich alkaline basaltsalways prevail, but from amore dynamic environment— either wherea hotspot interacts with a spreading ridge such as Iceland (Fitton et al.,2003) or where an oceanic plateau forms due to strong plume flux inthe case of Ontong Java (Fitton and Godard, 2004).

5.3.2. Subduction component and mineral controlThough Nb/La has been used, Th/Nb is introduced here to help

distinguish Nb-anomalies of various origins since it shows no overlap

between oceanic mantle and continental crust (Albarede, 1998).Based on data from modern arc systems worldwide, Plank (2005)concluded that a high Th/La ratio is a unique signature of continentalcrust and cannot be created by subduction processes such asmetamorphism of slab sediments as Th and REE in metapelites arenot fractionated throughout the prograde path up to blueschist oreven ecolgite-facies (Arculus et al., 1999; Spandler et al., 2003). Thus,elevated Th/La in some arcmagmas can only be attributed to sedimentrecycling (Plank and Langmuir, 1993). As mixing with bulk sedimentfails to produce the trace element ratios and isotopic signatures of arcmagmas (Turner et al., 1996; Elliott et al., 1997; Class et al., 2000), Thenrichment in the subduction zone is probably caused by addition of

Fig. 10. a. Plot of Zr/Y−Nb/Y (ΔNb=0 reference line from Fitton et al., 1997; fieldsand data points from Condie, 2005). Abbreviations: ARC, arc-related basalts;NMORB, normal mid oceanic ridge basalt; OIB, oceanic island basalt; OPB, oceanicplateau basalt. b. Plot of La/Yb vs. Zr/Nb (after Aldanmaz et al., 2008; melt curvesobtained non-modal batch melting model by Aldanmaz et al., 2006 for details ofsource data of various mantle source compositions and MORB field). Numbers oncurves denote degree of partial melting; abbreviations: DM, depleted mantle; PM,primitive mantle; Spl, spinel; Grt, garnet. c. Zr/Y vs. Nb/Zr plots (please refer to fig. 9of Ichiyama et al., 2008 for the fields of representative data from oceanic islands,oceanic plateaux and sea-floor spreading axis, as well as the three partial meltcomposition curves modeled by modal batch melting equations).

Fig. 11. a. La/Nb vs. Th/Nb (after Marchesi et al. (2007); values of crustal componentsand data of various fields from Plank (2005)). b. Nd/Yb–Hf/Yb variation diagram (afterPearce et al., 1999). Displacement from mantle trend is denoted by ΔHf (+ve Hfanomaly) and ΔNd (−ve Hf anomaly). The curves are mixing trends between DM(depleted mantle) and subducted pelagic sediment modeled with a mass fraction from0.05 to 0.2 of subduction component in the mantle. Ratios for Hf and Nd betweensubduction component and mantle are represented by rHf and rNd. (Please refer to thereference paper for definitions and modeling method). c. Zr/Sm vs. Nb/La plot(modified from fig. 8 of Munker et al., 2004; MORB composition after Hofmann, 1988).Rutile, because of its high DZr,Nb/DREE, would cause a positive trend when it is in control.Inverse correlation, on the other hand, would be observed for the residual liquid if low-Mg amphibole with high DNb/DLa and low DZr/DSm is fractionated.


sediment melt (Johnson and Plank, 1999). As shown in the La/Nb vs.Th/Nb plot (Fig. 11a; after Marchesi et al., 2007), a few T-MORBsamples and all subgroup A members of the SSZ-type samples locatewell above the MORB-OIB array, indicating input of subduction-derived Th. Also noticeable are the two trends. Whereas the oneformed by T-MORBs (Th/La=0.1) extends towards the field ofMarianas arc basalts, that of subgroup A members (Th/La from 0.2to 0.3) overlaps the lower part of the field of Aleutians arc basalts.

The elemental hafnium–neodymium system also provides goodconstraints on inputs of subduction-derived material. During partialmelting of mantle under normal conditions, Hf and Nd generallymaintain coherent behaviors as they share similar incompatibilities.Depletion or enrichment of Hf with respect to Nd (ΔHf), however,would occur in a subduction environment when slab-derived fluid isadded (−ΔHf) or low-Mg amphibole is being fractionated (+ΔHf)


(Pearce et al., 1999). As shown in Fig. 11b, the T-MORB group and, to aless extent, the OIB-type samples show negative ΔHf trends. Thisprobably indicates the involvement of slab-derived fluids rather thanmelt(s) of sediment and/or oceanic crust since Hf has relatively lowsolubility in aqueous fluids (Barry et al., 2006). On the other hand, theSSZ-type samples show slightly positive ΔHf. Such enrichment of Hfrelative to REE is likely caused by fractional crystallization of low-Mgamphiboles which is supported by a negative correlation between Nb/La and Zr/Sm (Fig. 11c; after Munker et al., 2004). Since formation oflow-Mg amphibole may be facilitated by infiltration of water-richsilicic melts or a silica-rich aqueous fluid (Ionov and Hofmann, 1995;Laurora et al., 2001), thesemagmas probably originated from amantlewedge metasomatized by slab-derived material. Also shown in thisfigure is a positive correlation of the ratios of T-MORB samples whichseems to be caused by fractional crystallization of rutile. However,since rutile is highly soluble in basaltic melts and is thus not stableduring peridotite melting (Ryerson and Watson, 1987), it is verypossible that this trend reflects interaction between basalts andaqueous fluid generated from slab dehydration instead. In asubduction zone, rutile forms from ilmenite and/or titanite duringhigh-pressure metamorphism (Zack et al., 2002). This mineral, as aresidual phase, retains significant amounts of conservative elements(Nb, Ta, Zr and Hf) in the down-dragged slab (Brenan et al., 1995;Stalder et al., 1998), leaving the fluid released into the overlyingmantle depleted in HFSE relative to REE (McCulloch and Gamble,1991).

The above discussion unequivocally indicates that the formation ofAOC involved more than one type of subduction-related material.Clearly, mantle-derived melts, particularly the T-MORB samples, weresubject to influx of a slab-derived aqueous fluid. Yet, as suggested bythe large range in ΔHf values from zero to moderately negative, thedegree of such contamination is highly variable. In addition, limitedparticipation of recycled sediment in the T-MORB was revealed byslight thorium enrichment of a few members. In terms of Th/La, theseMORB samples demonstrate a great resemblance to basalts from theMariana arc of which the subducted sediments are clay, chert andvolcanoclastic rocks (Karpoff, 1992). Therefore, a certain amount ofpelagic sediment withminor arc-related material was probably addedto the source of T-MORB.

Unlike the other two groupswith onlyminor subduction influence,the SSZ-type samples are genetically linked to arc system. Theirpositive ΔHf suggest that they originated from mantle wedge whichwas metasomatized by a silica-and-water-rich phase. In spite of thesame source, the two subgroups of the SSZ-type samples differ fromeach other in the level of thorium enrichment. Whereas the subgroupB shows very limited input of sediment melt, subgroup A resemblebasalts from the Central and Eastern Aleutian arc (fig. 3 of Singer et al.,2007) of which terrigenous sediment is the dominant materialrecycled (Scholl and Creager, 1973).

5.4. Petrogenetic considerations

5.4.1. Origin of the MORB- and OIB-type samplesIt has long been thought that MORBs are chemically uniform

because they are products of decompressional melting of a homoge-neous reservoir on a global scale (Dupre et al., 1981; Macdougall andLugmair, 1985; Zindler and Hart, 1986). In this sense, enrichedmantlesources found in OIB and E-MORB are representation of minorheterogeneities in the depleted upper mantle (Sun et al., 1979;Kellogg and Turcotte, 1990). Often these enriched components areascribed to recycled material such as ancient crust, lithosphericfragments or sediments that were once dragged down to variousdepth of mantle during delamination or subduction and then broughtback to the surface by positive buoyancy and/or mantle convection orthey can be deep-seated plumes arising from the core–mantle

boundary (Allegre and Turcotte, 1985; Hart, 1988; Weaver, 1991;Hart et al., 1992; Hofmann, 1997; Keken et al., 2002; Anderson, 2007).

At first glance, our OIB-type and E-MORB samples may representan oceanic island such as Ascension and Tristan da Cunha in whichenriched basalt is always the dominant lithology (Wilson, 1989). Ifthis is the case, it implies that the AOC is a tectonic mélange — anexotic oceanic island juxtaposed with an ophiolite composed of T-MORB and SSZ-type rocks. However, not only our field data do notsupport this hypothesis, the slightly negative ΔHf (Fig. 11b) and thedepletion in Zr and Hf relative to Nd (Fig. 7a and b) for the OIB-typeand E-MORB samples indicates involvement of a slab-derived aqueousfluid which is rather unusual for a typical oceanic island. While theinfluence of a slab-derived fluid on MORB-type samples indicates thatboth types of samples were formed in the vicinity of a subductionzone, mixing trend (Fig. 10b) suggests that they interacted with eachother before emplacement.

As mentioned in last section, genesis of the AOC is linked to adynamic environment, e.g. oceanic plateau formation (Fitton andGodard, 2004) or interaction between a mid-ocean ridge and plume(Rhodes et al., 1990; Douglass et al., 1999). Despite the geochemicalaffinities, these two mechanisms are very unlikely. Firstly, the greatvariety in composition of our MORB-type samples is at odds with theremarkable geochemical uniformity that characterizes tholeiiticbasalts of the Ontong Java Plateau (Mahoney et al., 1993; Neal et al.,1997; Tejada et al., 2002). Secondly, the interaction of our MORB-typesamples with slab-derived fluid and pelagic sediments is a processrarely found in ocean island setting.

Recent discoveries from mid-ocean ridge systems worldwideindicate that enriched basalts are not only limited to hotspot orplume settings but are actually widespread and can be found in bothfast- and slow-spreading ridges (Niu et al., 1999; Hemond et al., 2006;Nauret et al., 2006). Modern data compilation also indicates thatmantle heterogeneity is far more common than previously thought(Kent et al., 2004; Janney et al., 2005; Graham et al., 2006; Fitton,2007), and there exists a complete continuum between depleted andenriched MORBs (Hofmann, 2003). As an alternative to recycled crustor lithosphere which are always needed for the plume model,metasomatic veins in lithospheric mantle can be the main source ofthe enrichment (Niu et al., 2002; Workman et al., 2004; Donnelly etal., 2004; Pilet et al., 2005). If so, the enriched mantle componentsfound in the AOC are hornblendite veins which were formed byinfiltration of low-degree melts of the Low Velocity Zone into oceaniclithosphere (Pilet et al., 2008; Niu, 2008).

Since some T-MORB and OIB-type samples bear a slab-derivedsignature, the origin of the AOC should be explained by a sea-floorspreading event involving interaction between a heterogeneousasthenospheric source and a subduction zone. Beneath an activemid-ocean ridge, decompression melting of blocks of ancientlithosphere which was previously metasomatized are widespread inthe upper mantle. Owing to their relatively lower solidus tempera-ture, the amphibolite veins in those recycled lithospheric fragmentsmelt before peridotite and produce OIB-type alkaline melts with aconspicuous Nb–Ta spike (Fig. 7a). During ascent, these melts mixwith the melts of the depleted mantle to produce the MORBs withvariable degrees of enrichment. These hybrid melts then come intocontact with a subduction zone and thus interact with slab-derivedfluid and pelagic sediments.

5.4.2. Origin of the SSZ-type samplesThe SSZ-type samples are depleted in HFSE relative to N-MORB

(Fig. 7c). Given the incompatibility of HFSE, no plausible degree of onesingle melting of a normal upper mantle source could yield suchcharacteristics. The origin can best be illustrated by the two-stagemelting model (Pearce, 1982; Woodhead et al., 1993; Hawkesworthet al., 1993; Pearce et al., 2005; Pearce and Robinson, 2010). To beginwith, there exists a mantle wedge source which is depleted in all


incompatible elements by prior melt extraction. Subsequent additionof melt and fluid derived from dehydration melting of the subductedsediments and the slab itself (Ringwood, 1977; Pearce and Peate,1995; Pearce et al., 1995) replenishes subduction-mobile elementsand helps trigger the melting of the mantle wedge by lowering thesolidus. This model implies that the more depleted a mantle source,the more susceptible its melt may display the effects of suchmetasomatism (Ewart and Hawkesworth, 1987). This model is veryexplanatory in a sense that the silicic melt and aqueous fluid added tothe mantle wedge will actually facilitate formation of low-Mgamphibole which, being fractionated (Fig. 11c), can lead to slightenrichment of Hf relative to Nd (Fig. 11b).

Though the above model agrees well with the general character-istics of the SSZ-type samples, differences between the two subgroupsneed to be addressed. The lower LREE contents of subgroup B samplescan neither be the result of a stronger depletion by first-stage meltingnor the consequence of a smaller degree of second-stage melting;otherwise, the behavior of other trace elements (fromNd, through theincompatible Zr and Ti, all the way to Lu) will not be strikingly similarto those of subgroup A samples (Fig. 7c). Because subgroup A samplesare more evolved (richer in silica; lower in MgOwt.% andMg number;lower in Ni content), it is possible that their elevated LREE content iscaused by fractionation of some phases commonly found in basalticsystems. Among these, only clinopyroxene (Cpx) and hornblende(Hbl; assumed to be low-Mg amphibole here) are able to enrich theresidual melt in LREE without modifying much the MREE and HREEpattern (fig. 4.8 of Rollision, 1993). Low-Mg amphibole, due to its highAmph/LDNb/Zr, incorporates Nb (and Ta) in amounts far greater than Zr(and Hf) (Tiepolo et al., 2001). Therefore, the larger the degree of itsfractionation, the lower the Nb/Zr of a residual melt. Since it is thesubgroup B ((Nb/Zr)N=0.558) and not subgroup A ((Nb/Zr)N=0.767) that is more depleted in Nb relative to Zr, we suggestthat the lower Nb/Zr of subgroup B is the result of removal of Hbl(low-Mg amphibole) to amore significant amountwhereas the higherLREE content of subgroup A is caused by a stronger fractionation ofCpx.

To reconcile the separate evolution paths of the two subgroupswith their contrasting Th contents, we have to turn to the very natureof subduction zone itself, i.e. production of different slab-derivedcomponents at various depths. According to Nichols et al. (1994) andTatsumi and Eggins (1995), dehydration melting of sediments andhydrousminerals is concentrated in several zones (see alsoMaruyamaet al., 2009; Nakamura and Iwamori, 2009). The first is wheretemperature rises to about 580 °C. In this region, serpentines of thesubducting slab break down and release water-rich fluid which isresponsible formetasomatism of themantlewedge.We postulate thatsubgroup B samples were originated from this zone for two reasons—1) the slab surface is not hot enough to melt the hydrated sedimentswhich is the repository of continent-derived Th; this explains their lowTh contents, and 2) the relatively low temperature (b900 °C) andpressure (b30 kbar) of the overlying mantle favor fractional crystal-lization of Hbl rather than Cpx and thus lower the Nb/Zr of a melt. Onthe other hand, both the enrichment of Th from recycled sediments(Fig. 11a) and the dominance of Cpx over Hbl as a phase controllingLREE fractionation tend to support that subgroupA samples originatedfrom the second zone which was deeper and further away from thetrench. In this zone, the temperature and pressure (∼650 °C and35 kbar) were high enough to trigger both decomposition ofamphiboles and dehydration melting of subducted sediments.

Taking also into account the affinities to a MORB-formingenvironment shown by the Nb–Zr systematics of several SSZ-typesamples (Fig. 10a and c), it is likely that the AOC was formed in aforearc spreading environment. Due to decompression, the near-trench metasomatized mantle wedge melted and produced thesubgroup B samples with MORB-like character and HFSE depletion.Due to the proximity, pre-existing arc-related melts from the deeper

region of the subduction zone may be drawn into the spreading basinand later become subgroup A samples. Whereas their higher Thconcentrations suggest that subgroup A samples originated from asource modified by sediment melts in addition to slab-derivedaqueous fluid, their elevated LREE and higher Nb/Zr are the resultsof stronger fractionation of Cpx than Hbl (low-Mg amphibole) whichis a combined effect of greater depth and higher temperature.

6. Discussion

6.1. Tectonic model

From the previous discussions, the origin of the AOC should beelucidated with a scenario which can cause intra-arc spreading whileallowing a mantle wedge to interact with an asthenospheric source.Obviously, ridge subduction is the best model (e.g., Santosh andKusky, 2010; Zhang et al., 2010). In many parts of the world such asAlaska (Sisson and Pavlis, 1993; Davis et al., 1998), Japan (Kinoshita,1995; Maruyama, 1997), and Latin America (Guivel et al., 2003; Cliftet al., 2003; Bellon et al., 2006), a spreading ridge interacts with asubduction zone and leaves mafic complexes or ophiolites withremarkably diverse compositions. Whereas tholeiites with N- and E-MORB signatures are ubiquitous in the area above the slab window,back-arc basin basalts and arc-related calc-alkaline basalts are notrare (Lytwyn and Casey, 1995; Guivel et al., 1999; Gutierrez et al.,2005). In fact, alkaline basalts that are geochemically indistinguish-able from ocean–island and/or plume-related basalts are alsocommonly found as they are closely associated with asthenosphericupwelling (Hole et al., 1991; Dostal et al., 2003; Espinoza et al., 2005).Very often, these mafic rock “cocktails” appear in the forearc or intra-arc region (Lytwyn et al., 1997; Shervais et al., 2005; Schoonmakerand Kidd, 2006) where the shortest possible course of emplacementspares the melts from complete homogenization. Overall, theproperties of AOC fit the ridge trench interaction model for SSZophiolite formation (Sisson et al., 2003; Pearce, 2003).

The last hurdle this model has to pass before it can be applied toAOC formation is the production of some special rock types which arecharacteristic for, if not unique to, ridge subduction settings. Adakite,as the product of anatexis of a hot subducted slab (Defant andDrummond, 1990; Thorkelson and Breitsprecher, 2005), often occursduring subduction of a spreading ridge (Johnston and Thorkelson,1997; Abratis and Worner, 2001; Percival et al., 2003; Breitsprecheret al., 2003; Zhang et al., 2010). High-Mg rocks including boninite andpicrite, on the other hand, represent melting of a refractory mantlewedge caused by the development of a slab window (McCarron andSmellie, 1998; Wilson, 2003). Sometimes, the slab melt interacts withdepleted peridotite and produces Nb-enriched basalts (NEB) (Augil-lon-Robles et al., 2001; Benoit et al., 2002). At the forearc region of theChinese Altai active margin (Terrane 6) to the southwest, adakites,Nb-enriched basalts, boninites and picrites are exposed extensively(Zhang et al., 2005a; Niu et al., 2006). Zircon U–Pb dating of theadakites yielded a concordant age of ca. 440 Ma (Zhang et al., 2008),indicating that these rocks are coeval with the Alegedayi OphiliticComplex.

We propose that at around 440 to 430 Ma, highly obliquesubduction of a spreading ridge began in the Chinese Altai(Fig. 12b). Formation of a slab window allowed influx of astheno-spheric mantle and thus caused the opening of an intra-arc basin.Although slab-derived fluids and pelagic sediments overprinted thisupwelling mantle with arc signatures, pre-subduction within-plateenrichment of this sub-slabmantlewas strong enough tomaintain theoriginal characteristics of the MORB and OIB-type samples. This sea-floor spreading over a supra-subduction zone also caused decom-pression melting in the mantle wedge and the encroachment of pre-existing melt from the arc which produced the two kinds of SSZ-typerocks. Because of the torch effect, the refractory peridotites at the

Fig. 12. Diagrams illustrating the tectonic evolution of the Chinese Altai from the Cambrian to the Early Permian. The evolution stages of the Aermantai intra-oceanic arc system arebased on the data from Li, 1991; Liu and Yuan, 1996; Han et al., 1997; He et al., 2001; Wang et al., 2003; Ping et al., 2005; Xiao et al., 2009. Please see the text for the source of dataused to reconstruct the tectonic history of the Chinese Altai.


forearc underwent partial melting and produced the boninites andpicrites. Meanwhile, at the two edges of the torn-apart ridge, theyoung ocean crust underwent partial fusion. Whereas melts from thelower edge reacted with the hot and thick ambient mantle duringascent and formed the Nb-enriched basalts, those from the upper edgepassed through the thin and cold peridotites without muchinteraction and were preserved as contemporary adakites.

6.2. Implications for the evolution of the Chinese Altai

From the Cambrian to Devonian, the Chinese Altai experiencedtwo tectonothermal events — one pre 450 Ma and one post 430 Ma(Tong et al., 2007). Although there is only a handful of direct examplesof the pre-450 Ma event— a 505 Ma rhyodacite (Windley et al., 2002),the 462 Ma Qiemuerqieke granitic gneiss (Wang et al., 2006), the U–Pb ages of many magmatic zircons retrieved from both gneissicgranitoids and meta- volcanoclastic rocks indicate that magmaticactivity prevailed from 540 to 450 Ma (Fig. 12a; Briggs et al., 2007;Long et al., 2007; Sun et al., 2008). Geochemical studies showed that—1) the only granitoid sample from this period has whole-rock εNd(from 0 to −1.2) (Wang et al., 2006), and all igneous zircons formedin this period have εHf(t) from −17.4 to +14.7 (Sun et al., 2008),indicating that both old crust and juvenile mantle were the source,2) Neoproterozoic inheritance can be found in zircons of both igneousand sedimentary rocks from this period, and 3) the geochemicalsignatures of all these rocks show a subduction origin. These resultsindicate that both old crust and juvenile mantle were involved in themagma source and suggest the Chinese Altai was a Cambrian toOrdovician arc built on a Neoproterozoic continental margin.

Post-430 Ma magmatism was extensive in the region. Numerousgranitic plutons with U–Pb zircon ages between 430 and 350 Mawerereported for Terranes 1 to 4 of the Chinese Altai (Lou, 1997; Zhanget al., 2000; Windley et al., 2002; Wang et al., 2006; Sun et al., 2008;Tong et al., 2007; Yuan et al., 2007). In addition, the 407 Ma bimodalvolcanic rocks of the Kangbutiebao Formation (Zhang et al., 2000), the405 Ma gabbro at Keketuohai (Wang et al., 2006), the 372 Ma Kuertimafic rocks (Zhang et al., 2003; Xu et al., 2003a,b), and the 352 MaBuergen ophiolitic belt (Wu et al., 2006) reveal that the entire areawas in an extensional setting. All available εNd(t) data from thisperiod are generally high (from −2.7 to +10.7) (Xu et al., 2003a,b;Wang et al., 2006; Niu et al., 2006; Yuan et al., 2007). Igneous zirconsformed in this period all show positive εHf(t) values from +2.3 to+17 (Sun et al., 2008). This implies that input from asthenosphericmantle was significant. Overall, the Late Silurian to Devonianevolution of the Chinese Altai can be regarded as large-scale anatexisof young arc crust, triggered by mantle upwelling.

Whereas it is clear that the tectonic setting before 450 Ma and after430 Ma was indeed fundamentally different, it is very surprising thatnot much attention has so far been paid to the cause of such dramaticchange. We believe that our discovery of Early Silurian ridgesubduction not only fills the apparent quiescence of about 20 Mabut also provides a more comprehensive perspective for an under-standing of this very abrupt and dramatic geodynamic transition. Asan active mid-ocean ridge approached the trench at about 450 Ma, arcmagmatism diminished because the supply of both sediment andwater dwindled with the decreasing age of the subducting slab. Whenthe spreading ridge finally became subducted at around 440 Ma(Fig. 12b), cessation of arc magmatism which is a common effect of


ridge–trench interaction (Cande and Leslie, 1986; Osozawa, 1997;Sisson et al., 2003) happened in the Chinese Altai and brought theCambrian to Devonian active margin setting to the end. Arc magmaticwas then replaced by intra-arc spreading which produced the AOCand by the occurrence of adakites, NEBs and picrites.

As the slab window continued to develop, the lower half of the slabwas finally detached from the upper part. The result was that theregion was no longer under the influence of slab pull and probablybecame an extensional regime (Fig. 12c). As extra space created by theregional extension kept inviting upwelling of the asthenosphere, theentire lower crust of the Chinese Altai underwent anatexis, resultingin the formation of voluminous granitoids. Besides, this could alsohave caused the magmatism and mineralization with prominentmantle signatures (Wang et al., 1999; Chen and Jahn, 2002; Chiaradiaet al., 2006; Wang et al., 2006; Yuan et al., 2007; Sun et al., 2008). It isalso likely that the large-scale emplacement of granitoid plutons ledto high-grade metamorphism in the country rocks and sediments (Xuet al., 2005; Wei et al., 2007) which peaked at around 390 to 380 Ma(Zheng et al., 2007; Long et al., 2007).

7. Conclusions

The results of our analyses show that the Alegedayi OphioliticComplex discovered in the Chinese Altai is a supra-subduction zoneophiolite with an enriched mantle signature emplaced at 439±17 Ma. It formed through mixing among various subduction-relatedmaterials and mantle components — a scenario which can best beexplained by interaction between a spreading ridge and a subductionzone. Considering other contemporary geological features, we favoran Early Silurian ridge subduction model for the origin of thisophiolitic complex because it explains not only the sudden cessationof arc magmatism which lasted from the Cambrian to the Ordovician,but also a rather significant mantle input found inmagmatic rocks andmineral deposits formed from the Late Silurian to the EarlyCarboniferous. In conclusion, the AOC recorded a tectonic eventwhich had profound effects on the evolution of the Chinese Altai.

Acknowledgements

This study was supported by Hong Kong RGC grants (HKU7043/07P), National Basic Research Program of China (2007CB411308),German/Hong Kong Joint Research Scheme sponsored by RGC of HongKong and GAES of Germany (G-HK030/07). We thank Yildirim Dilek,Hugh Rollinson, Alfred Kröner, Brian Windley, Bor-Ming Jahn andKent Condie for their invaluable comments.

References

Abratis, M., Worner, G., 2001. Ridge collision, slab-window formation, and the flux ofPacific asthenosphere into the Caribbean realm. Geology 29, 127–130.

Albarede, F., 1998. The growth of continental crust. Tectonophysics 296, 1–14.Aldanmaz, E., Yaliniz, M.K., Guctekin, A., Goncuoglu, M.C., 2008. Geochemical

charactistics of mafic lavas from the Neoththyan ophiolites in western Turkey:implications for heterogeneous source contribution during variable stages of oceancrust generation. Geological Magazine 145, 37–54.

Aldanmaz, E., Koprubasi, N., Gurer, Ö.F., Kaymakci, N., Gourgaud, A., 2006. Geochemicalconstraints on the Cenozoic, OIB-type alkaline volcanic rocks of NW Turkey:Implications for mantle sources and melting processes. Lithos 86, 50–76.

Allegre, C.J., Turcotte, D.L., 1985. Geodynamic mixing in the mesosphere boundary layerand the origin of oceanic islands. Geophysical Research Letters 12, 207–210.

Anderson, D.L., 2007. New Theory of the Earth. Cambridge University Press, New York.49–61.

Ao, S.J., Xiao, W.J., Han, C.M., Mao, Q.G., Zhang, J.E., 2010. Geochronology andgeochemistry of Early Permian mafic–ultramafic complexes in the Beishan area,Xinjiang, NW China: implications for late Paleozoic tectonic evolution of thesouthern Altaids. Gondwana Research 18, 466–478 (this issue).

Arculus, R.J., Lapierre, H., Jaillard, E., 1999. Geochemical window into subduction andaccretion processes: Raspas metamorphic complex, Ecuador. Geology 27, 547–550.

Augillon-Robles, A., Calmus, T., Benoit, M., Bellon, H., Maury, R.C., Cotton, J., Bourgois, J.,Michaud, F., 2001. Late Miocene adakites and Nb-enriched basalts from Vizcaino

Peninsula, Mexico: indicators of East Pacific Rise subduction below Baja California?Geology 29, 531–534.

Badarch, G., Cunningham, W.D., Windley, B.F., 2002. A new terrane subdivision forMongolia: implications for the Phanerozoic crustal growth of Central Asia. Journalof Asian Earth Sciences 21, 87–110.

Barry, T.L., Pearce, J.A., Leat, P.T., Millar, I.L., le Roex, A.P., 2006. Hf isotope evidence forselective mobility of high-field-strength elements in a subduction setting: SouthSandwich Islands. Earth and Planetary Science Letters 252, 223–244.

Bellon, H., Aguillon-Robles, A., Calmus, T., Maury, R.C., Bourgois, J., Cotton, J., 2006. LaPurisima volcanic field, Baja California Sur (Mexico): Miocene to Quaternaryvolcanism related to subduction and opening of an asthenospheric window. Journalof Volcanology and Geothermal Research 152, 253–272.

Benoit, M., Augillon-Robles, A., Calmus, T., Maury, R.C., Bellon, H., Cotton, J., Bourgois, J.,Michaud, F., 2002. Geochemical Diversity of Late Miocene Volcanism in SouthernBaja California, Mexico: implication of mantle and crustal sources during theopening of an asthenospheric window. Journal of Geology 110, 627–648.

Bienvenu, P., Bougault, H., Joron, M., Treuil, M., Dmitriev, L., 1990. MORB alteration:rare-earth element/ non-rare-earth hygromagmaphile element fractionation.Chemical Geology 82, 1–14.

Bourgois, J., Martin, H., Lagabrielle, Y., Le Moigne, J., Frutos Jara, J., 1996. Subductionerosion related to spreading-ridge subduction: Taitao peninsula (Chile margintriple junction area). Geology 24, 723–726.

Breitsprecher, K., Thorkelson, D.J., Groome, W.G., Dostal, J., 2003. Geochemicalconfirmation of the Kula–Farallon slab window beneath the Pacific Northwest inEocene time. Geology 31, 351–354.

Brenan, J.M., Shaw, H.F., Ryerson, F.J., Phinney, D.L., 1995. Mineral aqueous fluidpartitioning of trace elements at 900 °C and 2.0 GPa: constrains on the traceelement geochemistry of mantle and deep crustal fluids. Geochimica etCosmochimica Acta 59, 3331–3350.

Briggs, S.M., Yin, A., Manning, C.E., Chen, Z.L., Wang, X.F., Grove, M., 2007. Late Paleozoictectonic history of the Ertix fault in the Chinese Altai and its implications for thedevelopment of the Central Asian Orogenic System. Geological Society of AmericaBulletin 119, 944–960.

Buchan, C., Pfander, J., Kroner, A., Brewer, T.S., Tomurtogoo, O., Tamurhuu, D.,Cunningham, D., Windley, B.F., 2002. Timing of accretion and collisionaldeformation in the Central Asian Orogenic Belt: implications of granite geochro-nology in the Bayankhongor Ophiolite Zone. Chemical Geology 192, 23–45.

Buslov, M.M., Fujiwara, Y., Iwata, K., Semakov, N.N., 2004a. Late Paleozoic–EarlyMesozoic tectonics and geodynamics of Central Asia. Gondwana Research 7,791–808.

Buslov, M.M., Watanabe, T., Fujiwara, Y., Iwata, K., Smirnova, L.V., Yu, Safonova I.,Semakov, N.N., Kiryanova, A.P., 2004b. Late Paleozoic faults of the Altai region,Central Asia: tectonic pattern and model of formation. Journal of Asian EarthSciences 23, 655–671.

Cande, S.C., Leslie, R.B., 1986. Late Cenozoic tectonics of the southern Chile Trench.Journal of Geophysical Research 91, 471–496.

Chen, B., Jahn, B., 2002. Geochemical and isotopic studies of the sedimentary andgranitic rocks of the Altai orogen of northwest China and their tectonicimplications. Geological Magazine 139, 1–13.

Chiaradia, M., Konopelko, D., Seltmann, R., Cliff, R.A., 2006. Lead isotope variationsacross terrane boundaries of the Tien Shan and Chinese Altay. Mineralium Deposita41, 411–428.

Class, C., Miller, D.L., Goldstein, S.L., Langmuir, C.H., 2000. Distinguishing melt and fluidcomponents in Umnak Volcanics, Aleutian Arc. Geochemistry, Geophysics,Geosystems 1 1999GC000010.

Clift, P.D., Pecher, I., Kuukowshi, N., Hampel, A., 2003. Tectonic erosion of the Peruvianforearc, Lima Basin, by subduction and Nazca Ridge collision. Tectonics 22, 1023.doi:10.1029/2002TC001386.

Coleman, R.G., 1989. Continental growth of Northwest China. Tectonics 8, 621–635.Condie, K.C., 2005. High field strength element ratios in Archean basalts: a window to

evolving sources of mantle plumes? Lithos 79, 491–504.Daux, V., Crovisier, J.L., Hemond, C., Petit, J.C., 1994. Geochemical evolution of basaltic

rocks subjected to weathering: fate of the major elements, rare earth elements, andthorium. Geochimica et Cosmochimica Acta 58, 4941–4954.

Davis, J.S., Roeske, S.M., Karl, S.M., 1998. Late Cretaceous to Early Tertiary transtensiionand strain partitioning in the Chugach accretionary complex, SE Alaska. Journal ofStructural Geology 20, 639–654.

Defant, M.J., Drummond, M.S., 1990. Derivation of some modern arc magmas bymelting of young subducted lithosphere. Nature 347, 662–665.

Dilek, Y., 2003a. Ophiolite pulses, mantle plumes and orogeny. In: Dilek, Y., Robinson, P.T.(Eds.), Ophiolites in EarthHistory:Geological Society. London.Special Publications, 218.

Dilek, Y., 2003b. Ophiolite concept and its evolution. In: Dilek, Y., Newcomb, S. (Eds.),Concept and the Evolution of Geological Thought, Geological Society of AmericaSpecial Paper, 373.

Dobretsov, N.L., Buslov, M.M., 2004. Serpentinitic mélanges associated with HP and UHProcks in Central Asia. International Geology Reviews 46, 957–980.

Donnelly, K.E., Goldstein, S.L., Langmuir, C.H., Spieglman, M., 2004. Origin of enrichedocean ridge basalts and implications for mantle dynamics. Earth and PlanetaryScience Letters 226, 347–366.

Dostal, J., Breitsprecher, K., Church, B.N., Thorkelson, D., Hamilton, T.S., 2003. Eocenemelting of Precambrain lithospheric mantle: analcime-bearing volcanic rocks fromthe Challis–Kamloops belt of south central British Columbia. Journal of Volcanologyand Geothermal Research 126, 303–326.

Douglass, J., Schilling, J.G., Fontignie, D., 1999. Plume–ridge interactions of theDiscovery and Shona mantle plumes with the southern mid-Atlantic Ridge (40º–55ºS). Journal of Geophysical Research 104, 2941–2962.

http://dx.doi.org/10.1029/2002TC001386


Dupre, B., Lambert, B., Rousseau, D., Allegre, C.J., 1981. Limitation on the scale of mantleheterogeneity under oceanic ridges. Nature 294, 552–554.

Elliott, T., Plank, T., Zindler, A., White, W., Bourdon, B., 1997. Element transport fromsubducted slab to volcanic front at theMariana arc. Journal of Geophysical Research102, 14991–15019.

Espinoza, F.,Morata, D., Pelleter, E.,Maury, R.C., Suarez,M., Lagabrielle, Y., Polve,M., Bellon,H., Cotton, J., De la Cruz, R., Guivel, C., 2005. Petrogenesis of the Eocene and Mio-Pliocene alkaline basaltic magmatism in Meseta Chile Chico, southern Patagonia,Chile: evidence for the participation of two slab windows. Lithos 82, 315–343.

Ewart, A., Hawkesworth, C.J., 1987. The Pleistocene–recent Tonga–Kermadec arc lavas:interpretation of new isotopic and race earth data in terms of a depleted mantlesource model. Journal of Petrology 28, 495–530.

Faure, G., Mensing, T.M., 2005. Isotopes: Principles and Applications, 3rd ed. JohnWileyand Sons Press.

Fedorovskii, V.S., Khain, E.E., Vladimirov, A.G., Kargopolov, S.A., Gibsher, A.S., Izokh, A.E.,1995. Tectonics, metamorphism and magmatism of collisional zones of the CentralAsian Caledonides. Geotectonics 29, 193–212.

Fitton, J.G., 2007. The OIB paradox. In: Foulger, G.R., Jurdy, D.M. (Eds.), Plates, Plumesand Planetary Processes.

Fitton, J.G., Godard, M., 2004. Origin and evolution of magmas on the Ontong JavaPlateau. In: Fitton, J.G., Mahoney, J.J., Wallace, P.J., Saunders, A.D. (Eds.), Origin andEvolution of the Ontong Java Plateau: Geological Society, London, SpecialPublications, 229, pp. 151–178.

Fitton, J.G., Saunders, A.D., Norry, M.J., Hardarson, B.S., Taylor, R.N., 1997. Thermal andchemical structure of the Iceland plume. Earth and Planetary Science Letters 153,197–208.

Fitton, J.G., Saunders, A.D., Kempton, P.D., Hardarson, B.S., 2003. Does depleted mantleform an intrinsic part of the Iceland plume? Geochemistry, Geophysics, Geosystems4. doi:10.1029/2002GC000424.

Floyd, P.A., Winchester, J.A., 1975. Magma-type and tectonic setting discriminationusing immobile elements. Earth and Planetary Science Letters 27, 211–218.

Graham, D.W., Elichert-Toft, J., Russo, J.C., Rubin, K.H., Albarede, F., 2006. Crypticstriations in the upper mantle revealed by hafnium isotopes in southeast Indianridges basalts. Nature 440, 199–202.

Gribble, R.F., Stern, R.J., Bloomer, S.H., Stuben, D., O'Hearn, T., Newman, S., 1996. MORBmantle and subduction components interact to generate basalts in the southernMariana Trough back-arc basin. Geochimica et Cosmochimica Acta 60, 2153–2166.

Gribble, R.F., Stern, R.J., Newman, S., Bloomer, S.H., O'Hearn, T., 1998. Chemical andisotopic composition of lavas from the Northern Mariana Trough: implications formagmagenesis in back-arc basins. Journal of Petrology 39, 125–154.

Guivel, C., Lagabrielle, Y., Bourgois, J., Maury, R.C., Fourcade, S., Martin, H., Arnaud, N.,1999. New geochemical constraints for the origin of ridge-subduction-relatedplutonic and volcanic suties from the Chile Triple Junction (Taitao Penninsula andSite 862, LEG ODP141 on the Taitao Ridge). Tectonophysics 311, 83–111.

Guivel, C., Lagabrielle, Y., Bourgois, J., Martin, H., Arnaud, N., Fourcade, S., Cotton, J.,Maury, R.C., 2003. Very shallow melting of oceanic crust during spreading ridgesubduction: origin of near-trench Quaternary volcanism at the Chile TripleJunction. Journal of Geophysical Research 108. doi:10.1029/2002JB002119.

Gutierrez, F., Gioncada, A., Gonzalez, Ferran O., Lashsen, A., Mazzuoli, R., 2005. TheHudson Volcano and surrounding monogenetic centres (Chilean Patagonia): anexample of volcanism associated with ridge–trench collision environment. Journalof Volcanology and Geothermal Research 145, 207–233.

Han, B., Wang, S., Jahn, B., Hong, D., Kagami, H., Sun, Y., 1997. Depleted-mantle sourcefor the Ulungur River A-type granites from North Xinjiang, China: geochemistryand Nd–Sr isotopic evidence, and implications for Phanerozoic crustal growth.Chemical Geology 138, 135–159.

Hart, S.R., 1988. Heterogeneous mantle domains: signatures, genesis and mixingchronologies. Earth and Planetary Science Letters 90, 273–296.

Hart, S.R., Hauri, E.H., Oschmann, L.A., Whitehead, J.A., 1992. Mantle plumes andentrainment: isotopic evidence. Science 256, 517–520.

Hartmann, L.A., Santos, J.O.S., 2004. Predominance of high Th/U magmatic zircons inBrazilian shield sandstones. Geology 32, 73–76.

Hawkesworth, C.J., Gallagher, K., Hergt, J.M., McDermott, F., 1993. Trace elementfractionation processes in the generation of island arc basalts. PhilosophicalTransactions. Royal Society of London A342, 179–191.

He, G., Han, B., Yue, Y., Wang, J., 1990. Tectonic division and crustal evolution of Altayorogenic belt in China. Geoscience Xinjiang 2, 9–20 (in Chinese with Englishabstract).

He, G.Q., Li, J.Y., Hao, J., Li, J.L., Cheng, S.D., Xu, X., Xiao, X.C., Tian, P.R., Deng, Z.Q., Li, Y.A.,Guo, F.X., 2001. Crustal Structure and Evolution of Xinjiang, China, Chinese National305 Project 07-01. Urumqi, Xinjiang, China.

Helo, C., Hegner, E., Kroner, A., Badarch, G., Tomurtogoo, O., Windley, B.F., Dulski, P.,2006. Geochemial signature of Paleozoic accretionary complexes of the CentralAsian Orogenic Belt in South Mongolia: constraints on arc environments andcrustal growth. Chemical Geology 227, 236–257.

Hemond, C., Hofmann, A.W., Vlastelic, I., Nauret, F., 2006. Origin of MORB enrichmentand relative trace element compatibilities along theMid-Atlantic Ridge between 10ºand 24ºN. Geochemistry, Geophysics, Geosystems 7. doi:10.1029/2006GC001317.

Hofmann, A.W., 1988. Chemical differentiation of the earth: the relationship betweenmantle, continental crust and oceanic crust. Earth and Planetary Science Letters 90,297–314.

Hofmann, A.W., 1997. Mantle geochemistry: the message from oceanic volcanism.Nature 385, 219–229.

Hofmann, A.W., 2003. Sampling mantle heterogeneity through oceanic basalts:isotopes and trace elements. In: Carlson, R.W. (Ed.), Treatise on Geochemistry.The Mantle and Core, 2. Elsevier, New York, pp. 61–101.

Hole, M.J., Rogers, G., Saunders, A.D., Storey, M., 1991. Relation between alkalicvolcanism and slab-window formation. Geology 19, 657–660.

Hoskin, P.W.O., Black, L.P., 2000. Metamorphic zircon formation by solid-staterecrystallization of protolith igneous zircons. Journal of Metamorphic Geology 18,423–439.

Hu, A., Jahn, B., Zhang, G., Chen, Y., Zhang, Q., 2000. Crustal evolution and Phanerozoiccrustal growth in northern Xinjiang: Nd isotopic evidence. Part I. Isotopiccharacterization of basement rocks. Tectonophysics 328, 15–51.

Hu, A., Zhang, G., Chen, Y., et al., 2006. Isotope Geochronology and Geochemistryfor Major Geological Events of Continental Crustal Evolution of Xinjiang, China,pp. 162–183.

Ichiyama, Y., Ishiwatari, A., Koizumi, K., 2008. Petrogenesis of greenstones from theMino-Tamba belt, SW Japan: evidence for an accreted Permian oceanic plateau.Lithos 100, 127–146.

Ionov, D.A., Hofmann, A.W., 1995. Nb–Ta-rich mantle amphiboles and micas:implications for subduction-related metasomatic trace element fractionations.Earth and Planetary Science Letters 131, 341–356.

Jahn, B.M., 2004. The Central Asian Orogenic Belt and growth of the continental crust inthe Phanerozoic. In: Malpas, J., Fletcher, C.J.N., Ali, J.R., Aitchison, J.C. (Eds.), Aspectsof the Tectonic Evolution of China: Geological Society, London, Special Publications,226, pp. 73–100.

Jahn, B.M., Wu, F., Chen, B., 2000. Granitoids of the Central Asian Orogenic Belt andcontinental growth in the Phanerozoic. Transactions of the Royal Society ofEdinburgh: Earth Sciences 91, 181–193.

Janney, P.E., Roex, A.P.L., Carlson, R.W., 2005. Hafnium isotope and trace elementconstraints on the nature of mantle heterogeneity beneath the Central SourthwestIndian Ridge (13ºE to 47ºE). Journal of Petrology 46, 2427–2464.

Johnson, M.C., Plank, T., 1999. Dehydration and melting experiments constrain the fateof subducted sediments. Geochemistry, Geophysics and Geosystems 11999GC000014.

Johnston, S.T., Thorkelson, D.J., 1997. Cocos–Nazca slab window beneath CentralAmerica. Earth and Planetary Science Letters 146, 465–474.

Karpoff, A.M., 1992. Cenozoic and Mesozoic sediments from the pigafetta basin, LEG129, sites 800 and 801: mineralogical and geochemical trends of the depositsoverlying the oldest oceanic crust. In: Larson, R.L., Lancelot, Y., et al. (Eds.),Proceedings of ODP, Scientific Results, 129. Ocean Drilling Program, College Station,TX, pp. 3–30.

Keken, P.E.v., Hauri, E., Ballentine, C.J., 2002. Mantle mixing: the generation,preservation, and destruction of chemical heterogeneity. Annual Review of Earthand Planetary Sciences 30, 493–525.

Kellogg, L., Turcotte, D., 1990. Mixing and the distribution of heterogeneities in achaotically convecting mantle. Journal of Geophysical Research 95, 421–432.

Kent, A.J.R., Stolper, E.M., Francis, D., Woodhead, J., Frei, R., Eiler, J., 2004. Mantleheterogeneity during the formation of the North Atlantic Igneous Province:constraints from trace element and Sr–Nd–Os–O isotope systemtics of Baffin Islandpicrites. Geochemistry Geophysics Geosystems 5. doi:10.1029/2004GC000743.

Kinoshita, O., 1995. Migration of igneous activities related to ridge subduction inSouthwest Japan and the East Asian continental margin from the Mesozoic to thePaleogene. Tectonophysics 245, 25–35.

Kovalenko, V.I., Yarmolyuk, V.V., Kovach, V.P., Kotov, A.B., Kozakov, I.K., Salnikova, E.B.,Larin, A.M., 2004. Isotopic provinces, mechanism of generation and sources of thecontinental crust in the Central Asian mobile belt: geological and isotopic evidence.Journal of Asian Earth Sciences 23, 605–627.

Kroner, A., Hegner, E., Lehmann, B., Heinhorst, J., Wingate, M.T.D., Liu, D.Y., Ermelov, P.,2008. Palaeozoic arc magmatism in the Central Asian Orogenic Belt of Kazakhstan:SHRIMP zircon ages and whole-rock Nd isotopic systematics. Journal of Asian EarthSciences 32, 118–130.

Lagabrielle, Y., Guivel, C., Maury, R., Bourgois, J., Fourcade, S., Martin, H., 2000.Magmatic–tectonic effects of high thermal regime at the site of active ridgesubduction: the Chile Triple Junction model. Tectonophysics 326, 255–268.

Laurent-Charvet, S., Charvet, J., 2003. Late Paleozoic strike-slip shear zones in easterncentral Asia (NW China): new structural and geochronological data. Tectonics 22,1009. doi:10.1029/2001TC901047.

Laurora, A., Mazzucchelli, M., Rivalenti, G., Vannucci, R., Zanetti, A., Barbieri, M.A.,Cingolani, C.A., 2001. Metasomatism and melting in carbonated peridotitexenoliths from the mantle wedge: the Gobernador Gregores case (SouthPetagonia). Journal of Petrology 42, 69–87.

Le Maitre, R.W., Bateman, P., Dudek, A., Keller, J., Lameyre, J., Le Bas, M.J., Sabine, P.A.,Schmid, R., Sorensen, H., Streckeisen, A., Woolley, A.R., Zanettin, B., 1989. AClassification of Igneous Rocks and Glossary of Terms. Blackwell, Oxford.

Li, J.Y., 1991. On evolution of Paleozoic plate tectonics of east Junggar, Xinjiang, China.In: Xiao, X.C., Tang, Y.Q. (Eds.), On Tectonic Evolution of the Southern Margin of thePaleozoic Composite Megasuture Zone. Beijing Technical Press, Beijing, pp. 92–108.

Liu, J.Y., Yuan, K.R., 1996. A discussion on the genesis and tectonic setting of alkaligranites in the Ulungur alkali-rich granite belt, Xinjiang. Geological Journal of ChinaUniversities 3, 257–272 (in Chinese with English abstract).

Long, X., Sun, M., Yuan, C., Xiao, W., Lin, S., Wu, F., Xia, X., Cai, K., 2007. Detrital zirconage and Hf isotopic studies for metasedimentary rocks from the Chinese Altai:implications for the Early Paleozoic tectonic evolution of the Central Asian OrogenicBelt. Tectonics 26. doi:10.1029/ 2007TC002128.

Lou, F., 1997. Characteristics of late Caledonian granite in the Nuoerte area Altay. JiangxiGeology 3, 60–66.

Ludden, J.N., Thompson, G., 1978. Behavior of rare earth elements during submarineweathering of tholeiitic basalts. Nature 274, 147–149.

Ludwig, K.R., 2001a. SQUID version 1.02: a geochronological toolkit for Microsoft Excel.Berkley Geochronological Centre Special Publication, 2, pp. 1–19.

http://dx.doi.org/10.1029/2002GC000424

http://dx.doi.org/10.1029/2002JB002119

http://dx.doi.org/10.1029/2006GC001317

http://dx.doi.org/10.1029/2004GC000743

http://dx.doi.org/10.1029/2001TC901047

http://dx.doi.org/10.1029/ 2007TC002128


Ludwig, K.R., 2001b. Users manual for Isoplot/ Ex. rev. 2.49. Berkley GeochronologicalCentre Special Publication, 1a, pp. 1–56.

Lytwyn, J.N., Casey, J.F., 1995. The geochemistry of postkinematic mafic dike swarmsand subophiolitic metabasites, Pozanti-karsanti ophiolite, Turkey: evidence forridge subduction. Geological Society of America Bulletin 107, 830–850.

Lytwyn, J., Casey, J., Gilbert, S., 1997. Arc-like mid-ocean ridge basalt formed seaward ofa trench–forearc system just prior to ridge subduction: an example fromsubaccreted ophiolites in southern Alaska. Journal of Geophysical Research 102,10225–10243.

Macdougall, J.D., Lugmair, G.W., 1985. Extreme isotopic homogeneity among basaltsfron the southern East Pacific Rise: mantle or mixing effect? Nature 313,209–211.

Mahoney, J.J., Storey, M., Duncan, R.A., Spencer, K.J., Pringle, M., 1993. Geochemistry andgeochronology of the Ontong Java Plateau. In: Pringle, M., Sager, W., Sliter, W.,Stein, S. (Eds.), The Mesozoic Pacific. Geology, Tectonics and Volcanism: AmericanGeophysical Union, Geophysical Monograph, 77, pp. 233–261.

Marchesi, C., Garrido, C.J., Bosch, D., Proenza, J.A., Gervilla, F., Monie, P., Rodriguez-Vega,A., 2007. Geochemistry of Cretaceous magmatism in eastern Cuba: recycling ofNorth American continental sediments and implications for subduction polarity inthe Greater Antilles paleo-arc. Journal of Petrology 48, 1813–1840.

Maruyama, S., 1997. Pacific-type orogeny revisited: Miyashiro-type orogeny proposed.Island Arc 6, 91–120.

Maruyama, S., Hasegawa, A., Santosh, M., Kogiso, T., Omori, S., Nakamura, H., Kawai, K.,Zhao, D., 2009. The dynamics of big mantle wedge, magma factory, andmetamorphic–metasomatic factory in subduction zones. Gondwana Research 16,414–430.

McCarron, J.J., Smellie, J.L., 1998. Tectonic implications of fore-arc magmatism andgeneration of high-magnesian andesites: Alexander Island, Antarctica. Journal ofthe Geological Society 155, 269–280.

McCulloch, M.T., Gamble, J.A., 1991. Geochemical and geodynamical constraints onsubduction zone magmatism. Earth and Planetary Science Letters 102, 358–374.

Miyashiro, A., 1974. Volcanic rock series in island arcs and active continental margins.American Journal of Science 274, 321–355.

Munker, C., Worner, G., Yogodzinski, G., Churikova, T., 2004. Behavior of high fieldstrength elements in subduction zones: constraints from Kamchatka–Aleutian arclavas. Earth and Planetary Science Letters 224, 275–293.

Nakamura, H., Iwamori, H., 2009. Contribution of slab-derived fluid in arc magmasbeneath the Japan arcs. Gondwana Research 16, 431–445.

Nauret, F., Abouchami, W., Galer, S.J.G., Hofmann, A.W., Hemond, C., Chauvel, C.,Dyment, J., 2006. Correlated trace element-Pb isotope enrichments in Indian MORBalong 10º–20ºS, Central Indian Ridge. Earth and Planetary Science Letters 245,137–152.

Neal, C.R., Mahoney, J.J., Kroenke, L.W., Duncan, R.A., Petterson, M.G., 1997. The OntongJava Plateau. In: Mahoney, J.J., Coffin, M.F. (Eds.), Large Igneous Provinces:Continental, Oceanic and Planetary Flood Volcanism. American GeophysicalUnion, Washington, D.C., pp. 183–216.

Nichols, G.T., Wyllie, P.J., Stern, C.R., 1994. Subduction zone melting of pelagicsediments constrained by melting experiments. Nature 371, 785–788.

Niu, Y.L., 2008. The origin of alkaline lavas. Science 320, 883–884.Niu, Y., Collerson, K.D., Batiza, R., 1999. Origin of enriched-type mid-ocean ridge basalt

at ridges far from mantle plumes: the East Pacific Rise at 11º20N. Journal ofGeophysical Research 104, 7067–7087.

Niu, Y., Regelous, M., Wendt, I.J., Batiza, R., O'Hara, M.J., 2002. Geochemistry of near-EPRseamounts: importance of source vs. process and the origin of enriched mantlecomponent. Earth and Planetary Science Letters 199, 327–345.

Niu, H., Sato, H., Zhang, H., Ito, J., Yu, X., Nago, T., Terada, K., Zhang, Q., 2006.Juxtaposition of adakite, boninite, high-TiO2 and low-TiO2 basalts in the Devoniansouthern Altay, Xinjiang, NW China. Journal of Asian Earth Science 28, 439–456.

Norrish, K., Hutton, J.T., 1969. An accurate X-ray spectrographic method for the analysisof a wide range of geological samples. Geochimica et Cosmochimica Acta 33,431–453.

Osozawa, S., 1997. The cessation of igneous activity and uplift when an activelyspreading ridge is subducted beneath an island arc. The Island Arc 6, 361–371.

Pearce, J.A., 1982. Trace element characteristics of lavas from destructive plate boundaries.In: Thorpe, R.S. (Ed.), Andesites. John Wiley and Sons, New York, pp. 525–548.

Pearce, J.A., 2003. Supra-subduction zone ophiolites: the search for modern analogues.In: Dilek, Y., Newcomb, S. (Eds.), Concept and the Evolution of Geological Thought,Geological Society of America Special Paper, p. 373.

Pearce, J.A., Peate, D.W., 1995. Tectonic implications of the composition of volcanic arcmagmas. Annual Review of Earth Planetary Sciences 23, 251–285.

Pearce, J.A., Robinson, P.T., 2010. The Troodos ophiolitic complex probably formed ina subduction initiation, slab edge setting. Gondwana Research 18, 60–81.

Pearce, J.A., Baker, P.E., Harvey, P.K., Luff, I.W., 1995. Geochemical evidence forsubduction fluxes, mantle melting and fractional crystallization beneath the SouthSandwich island arc. Journal of Petrology 36, 1073–1109.

Pearce, J.A., Kempton, P.D., Nowell, G.M., Noble, S.R., 1999. Hf–Nd element andisotope perspective on the nature and provenance of mantle and subductioncomponents in Western Pacific arc-basin systems. Journal of Petrology 40,1579–1611.

Pearce, J.A., Stern, R.J., Bloomer, S.H., Fryer, P., 2005. Geochemical mapping of theMariana arc–basin system: implications for the nature and distribution of thesubduction component. Geochemistry, Geophysics and Geosystems 6. doi:10.1029/2004GC000895.

Percival, J.A., Stern, R.A., Rayner, N., 2003. Archean adakites from the Ashuanipicomplex, eastern Superior Province, Canada: geochemistry, geochronology andtectonic significance. Contribution to Mineralogy and Petrology 145, 265–280.

Pilet, S., Hernandez, J., Sylvester, P., Poujol, M., 2005. The metasomatic alternative forocean island basalt chemical heterogeneity. Earth and Planetary Science Letters236, 148–166.

Pilet, S., Baker, M., Stolper, E.M., 2008. Metasomatized lithosphere and the origin ofalkaline lavas. Science 320, 917–919.

Ping, J., Dunyi, L., Yuruo, S., Fuqin, Z., 2005. SHRIMP dating of SSZ ophiolites fromnorthern Xinjiang Province, China: implications for generation of oceanic crust inthe Central Asian Orogenic Belt. In: Sklyarov, E.V. (Ed.), Structural and TectonicCorrelation across the Central Asia Orogenic Collage: North-Eastern Segment,Guidebook and Abstract Volume of the Siberian Workshop IGCP-480, IEC SBRAS,Irkutsk, p. 246.

Plank, T., 2005. Constraints from Thorium/ Lanthanum on sediment recycling atsubduction zones and the evolution of the Continents. Journal of Petrology 46,921–944.

Plank, T., Langmuir, C.H., 1993. Tracing trace elements from sediment input to volcanicoutput at subduction zones. Nature 362, 739–743.

Qi, L., Hu, J., Conrad, G.D., 2000. Determination of trace elements in granites byinductively coupled plasma mass spectrometry. Talanta 51, 507–513.

Qu, G., Chong, M., 1991. Lead isotope geology and its tectonic implications in theAltaids. China Geosciences 5, 100–110.

Rhodes, J.M., Morgan, C., Liias, R.A., 1990. Geochemistry of axial seamounts lavas:magmatic relationship between the Cobb hotspot and the Juan de Fuca ridge.Journal of Geophysical Research 95, 12713–12733.

Ringwood, A.E., 1977. Petrogenesis in island arc systems. In: Talwani, M., Pitman, I.W.C.(Eds.), Island Arcs Deep Sea Trenches and Back-arc Basins. Maurice Ewing Series 1.American Geophysical Union, Washington, D.C., pp. 311–324.

Rollision, H.R., 1993. Using Geochemical Data: Evaluation, Presentation, Interpretation.Longman Group UK Limited, Singapore.

Roser, B.P., Nathan, S., 1997. An evaluation of elemental mobility during metamorphismof a turbidite sequence (Greenland Group, New Zealand). Geological Magazine 134,219–234.

Ryerson, F.J., Watson, E.B., 1987. Rutile saturation in magmas: implications for Ti–Nb–Ta depletion in island arc basalts. Earth and Planetary Science Letters 86,225–239.

Safonova, I.Y., Utsunomiya, A., Kojima, S., Nakae, S., Tomurtogoo, O., Filippov, A.N.,Koizumi, K., 2009. Pacific superplume-related oceanic basalts hosted by accretion-ary complexes of Central Asia, Russian Far East and Japan. Gondwana Research 16,587–608.

Santosh, M., Kusky, T., 2010. Origin of paired high pressure–ultrahigh-temperatureorogens: a ridge subduction and slab window model. Terra Nova 22, 35–42.

Saunders, A.D., Tarney, J., 1984. Geochemical characteristics of basaltic volcanismwithin back-arc basins. In: Kokelaar, B.P., Howells, M.F. (Eds.), Marginal BasinGeology: Geological Society, London, Special Publications, 16, pp. 59–76.

Scholl, D.W., Creager, J.S., 1973. Geologic synthesis of Leg 19 (DSDP) results: far NorthPacific, and Aleutian Ridge, and Bering Sea. Deep Sea Drilling Projects. doi:10.2973/dsdp.proc.19.137.1973.

Schoonmaker, A., Kidd, W.S.F., 2006. Evidence for a ridge subduction event in theOrdovician rocks of north-central Maine. Geological Society of America Bulletin118, 897–912.

Sengor, A.C., Natal'in, B.A., 1996. Paleotectonics of Asia: fragments of a synthesis. In: Yin,A., Harrsion, M. (Eds.), The Tectonic Evolution of Asia. Cambridge University Press,Cambridge, pp. 486–640.

Sengor, A.C., Natal'in, B.A., Burtman, V.S., 1993. Evolution of the Altaid tectonic collageand Palaeozoic crustal growth in Eurasia. Nature 364, 299–306.

Shervais, J.W., Zoglman Schuman, M.M., Hanan, B.B., 2005. The Stonyford VolcanicComplex: a forearc seamount in the Northern California Coast Ranges. Journal ofPetrology 46, 2091–2128.

Shinjo, R., Chung, S.L., Kato, Y., Kimura, M., 1999. Geochemical and Sr–Nd isotopiccharacteristics of volcanic rocks from the Okinawa Trough and Ryukyu arc:implications for the evolution of a young, intracontinental back arc basin. Journal ofGeophysical Research 104, 10591–10608.

Singer, B.S., Jicha, b.R., Leeman, W.P., Rogers, N.W., Thirlwall, M.F., Ryan, J., Nicolaysen,K.E., 2007. Along-strike trace element and isotopic variation in Aleutian Island arcbasalt: subduction melts sediments and dehydrates serpentine. Journal ofGeophysical Research 112. doi:10.1029/ 2006JB004897.

Sisson, V.B., Pavlis, T.L., 1993. Geologic consequences of plate reorganization: anexample from the Eocene southern Alaska fore arc. Geology 21, 913–916.

Sisson, V.B., Pavlis, T.L., Roeske, S.M., Thorkelson, D.J., 2003. Introduction: an overviewof ridge–trench interactions in modern and ancient settings. In: Sisson, V.B.,Roeske, S.M., Pavlis, T.L. (Eds.), Geology of a Transpressional Orogen Developedduring Ridge–Trench Interaction along the North Pacific Margin: Geological Societyof America Special Paper, 371.

Spandler, C., Hermann, J., Arculus, R., Mavrogenes, J., 2003. Redistribution of traceelements during prograde metamorphism from lawsonite blueschist to eclogitefacies: implications for deep subduction-zone processes. Contributions to Miner-alogy and Petrology 146, 205–222.

Stalder, R., Foley, S.F., Brey, G.P., Horn, I., 1998. Mineral–aqueous fluid partitioning oftrace elements at 900–1200 ºC and 3.0–5.7 GPa: new experimental data for garnet,clinopyroxene and rutile, and implications for mantle metamosmatism. Geochi-mica et Cosomochimic Acta 62, 1781–1801.

Steiger, R.H., Jager, E., 1977. Subcommission on geochronology: convention on the useof decay constants in geo- and cosmochronology. Earth and Planetary ScienceLetters 36, 359–362.

Sun, S.S., McDonough, W.F., 1989. Chemcial and isotopic systematics and processes. In:Saunder, A.D., Norry, M.S. (Eds.), Magmatism in the Ocean Basins: GeologicalSociety of London, Special Publication, 42, pp. 313–345.

http://dx.doi.org/10.1029/2004GC000895

http://dx.doi.org/10.1029/2004GC000895

http://dx.doi.org/10.2973/dsdp.proc.19.137.1973

http://dx.doi.org/10.2973/dsdp.proc.19.137.1973

http://dx.doi.org/10.1029/ 2006JB004897


Sun, S.S., Nesbitt, R.W., Sharaskin, A.Y., 1979. Geochemical characteristics of mid-oceanridge basalts. Earth and Planetary Science Letters 44, 119–138.

Sun, M., Yuan, C., Xiao, W., Long, X., Xia, X., Zhao, G., Lin, S., Wu, F., Kroner, A., 2008.Zircon U–Pb and Hf isotopic study of gneissic rocks from the Chinese Altai:progressive accretionary history in the early to middle Paleozoic. Chemical Geology247, 352–383.

Sun, M., Long, X., Cai, K., Jiang, Y., Wang, B.Y., Yuan, C., Zhao, G., Xiao, W., Wu, F., 2009.Early Paleozoic ridge subduction in the Chinese Altai: insight from the abruptchange in zircon Hf isotopic compositions. Science in China, Series D 52, no.9,1345–1358.

Tatsumi, Y., Eggins, S., 1995. Subduction ZoneMagmatism. Blackwell, Oxford, pp. 69–74.Tatsumi, Y., Hamilton, D.L., Nesbitt, R.W., 1986. Chemical characteristics of fluid phase

from a subducted lithosphere and origin of arc magmas: evidence from high-pressure experiments and natural rocks. Journal of Volcanlogy and GeothermalResearch 29, 293–309.

Tejada, M.L.G., Mahoney, J.J., Neal, C.R., Duncan, R.A., Petterson, M.G., 2002. Basementgeochemisty and geochronology of Central Malaita, Solomon Islands, withimplications for the origin and evolution of the Ontong Java Plateau. Journal ofPetrology 43, 449–484.

Thirlwall, M.F., Smith, T.E., Graham, A.M., Theodorou, N., Hollings, P., Davidson, J.P.,Arculus, R.J., 1994. High field strength element anomalies in arc lavas: source orprocess? Journal of Petrology 35, 819–838.

Thorkelson, D.J., Breitsprecher, K., 2005. Partial melting of slab window margins:genesis of adakitic and non-adakitic magmas. Lithos 79, 25–41.

Tiepolo, M., Bottazzi, P., Foley, F., Oberti, R., Vannucci, R., Zanetti, A., 2001. Fractionationof Nb and Ta from Zr and Hf at mantle depths: the role of titanian pargasite andkaersutite. Journal of Petrology 42, 221–232.

Tong, Y., Wang, T., Hong, D.W., Han, B.F., Liu, X.M., 2007. Ages and origin of the earlyDevonian granites from the north part of Chinese Altai mountains and its tectonicimplications. Acta Petrologica Sinica 23, 1933–1944.

Turner, S.T., Hawkesworth, C.J., van Calsteren, P., Heath, E., Macdonald, R., Black, S.,1996. U-series isotopes and destructive margin magma genesis in the LesserAntilles. Earth and Planetary Science Letters 142, 191–207.

Wang, D., Chen, Y., Xu, Z., Li, H., 1999. Helium isotopic study on mantle degassing in theAltay orogenic zone, China. Chinese Science Bulletin 44, 1050–1053 (in Chinesewith English abstract).

Wang, Z., Sun, S., Li, J., Hou, Q., Qin, K., Xiao, W., Hao, J., 2003. Paleozoic tectonicevolution of the northern Xinjiang, China: geochemical and geochronologicalconstraints from the ophiolites. Tectonics 22. doi:10.1029/2002TC001396.

Wang, T., Hong, D., Tong, Y., Han, B., Shi, Y., 2005. Zircon U–Pb SHRIMP age and origin ofpost-orogenic Lamazhao granitic pluton from Altai orogen: its implications forvertical continental growth. Acta Petrologica Sinica 21, 640–650 (in Chinese withEnglish abstract).

Wang, T., Hong, D., Jahn, B., Tong, Y., Wang, Y., Han, B., Wang, X., 2006. Timing,petrogenesis and setting of Paleozoic synorogenic intrusions from the Altaimountains, northwest China: implications for the tectonic evolution of anaccretionary orogen. Journal of Geology 114, 735–751.

Wang, T., Jahn, B., Kovach, V., Tong, Y., Hong, D., Han, B., 2009. Nd–Sr isotopic mappingof the Chinese Altai and implications for continental growth in the Central AsianOrogenic Belt. Lithos 110, 359–372.

Weaver, B.L., 1991. Trace element evidence for the origin of ocean–island basalts.Geology 19, 123–126.

Wei, C., Clarke, G., Tian, W., Qiu, L., 2007. Transition of metamorphic series from thekyanite- to andalusite-types in the Altai orogen, Xinjiang, China: evidence frompetrography and calculated KMnFMASH and KFMASH phase relations. Lithos 96,353–374.

Williams, I.S., Claesson, S., 1987. Isotopic evidence for the Precambrian provenanceand Caledonian metamorphism of high grade paragneisses from the SeveNappes, Scandinavian Caledonides. Contribution to Mineralogy and Petrology97, 205–217.

Wilson, M., 1989. Igneous Petrogenesis. Unwin Hyman, London, pp. 245–285.Wilson, R.A., 2003. Geochemistry and petrogenesis of Ordovician arc-related mafic

volcanic rocks in the Popelogan Inlier, northern New Brunswick. Canadian Journalof Earth Sciences 40, 1171–1189.

Windley, B.F., Kroner, A., Guo, J., Qu, G., Li, Y., Zhang, C., 2002. Neoproterozoic toPaleozoic geology of the Altai Orogen, NW China: new zircon age data and tectonicevolution. Journal of Geology 110, 719–737.

Windley, B.F., Alexeiev, D., Xiao, W., Kroner, A., Badarch, G., 2007. Tectonic models foraccretion of the Central Asian Orogenic Belt. Journal of the Geological Society ofLondon 164, 31–47.

Wood, D.A., 1980. The application of a Th–Hf–Ta diagram to problems of tectono-magmatic classification and to establishing the nature of crustal contamination ofbasaltic lavas of the British Tertiary volcanic province. Earth and Planetary ScienceLetters 50, 11–30.

Woodhead, J., Eggins, S., Gamble, J., 1993. High field strength and transition elementsystematics in island arc and back-arc basalts: evidence for multi-phase melt

extraction and a depleted mantle wedge. Earth and Planetary Science Letters 114,491–504.

Workman, R.K., Hart, S.R., Jackson, M., Regelous, M., Farley, K.A., Blusztajn, J., Kurz, M.,Staudigel, H., 2004. Recycled metasomatized lithosphere as the origin of theEnriched Mantle II (EM2) end-member: evidence from the Samoan Volcanic Chain.Geochemisty, Geophysics, Geosystems 5. doi:10.1029/2003GC00623.

Wu, B., He, G.Q., Wu, T.R., Li, H.J., Luo, H.L., 2006. Discovery of the Buergen ophioliticmélange belt in Xinjiang and its tectonic significance. Geology in China 33, 476–486(in Chinese with English abstract).

Xiao, W., Kusky, T., 2009. Geodynamic processes and metallogenesis of the CentralAsian and related orogenic belts. Gondwana Research 16, 167–169.

Xiao, X.C., Tang, Y.Q., Feng, Y.M., Zhu, B.Q., Li, J.Y., Zhao, M., 1992. Tectonics in NorthernXinjiang and Its Neighboring Areas. Geological Publishing House, Beijing. (inChinese with English abstract).

Xiao, W., Windley, B.F., Badarch, G., Sun, S., Li, J., Qin, K., Wang, Z., 2004. Palaeozoicaccretionary and convergent tectonics of the southern Altaids: implications for thegrowth of Central Asia. Jounrnal of the Geological Society, London 161, 339–342.

Xiao,W.,Windley, B.F., Yan,Q., Qin,K., Chen,H., Yuan,C., Sun,M., Li, J., Sun, S., 2006. SHRIMPzircon age of the Aermentai ophiolite in the north Xinjiang area, China and its tectonicimplications. Acta Geologica Sinica 80, 32–37 (in Chinese with English abstract).

Xiao,W.,Windley, B.F., Yuan, C., Sun,M., Han, C.M., Lin, S.F., Chen, H.L., Yang, Q.R., Liu, D.Y.,Qin, K.Z., Li, J.L., Sun, S., 2009. Paleozoicmultiple subduction–accretion processes of thesouthern Altaids. American Journal of Science 309, 221–270.

Xiao, W., Huang, B., Han, C., Sun, S., Li, J., 2010. A review of the western part ofthe Altaids: a key to understanding the architecture of accretionary orogens.Gondwana Research 18, 253–273 (this issue).

Xu, J., Castillo, P.R., Chen, F., Niu, H., Yu, X., Zhen, Z., 2003a. Geochemistry of latePaleozoic mafic igneous rocks from the Kuerti area, Xinjiang, northewest China:implications for backarc mantle evolution. Chemical Geology 193, 137–154.

Xu, J., Chen, F., Yu, X., Niu, H., Zhen, Z., 2003b. Kuerti ophiolite in Altay area of northXinjiang: magmatism of an ancient back-arc basin. Acta Petrologica et Mineralogica20, 344–352 (in Chinese with English abstract).

Xu, X., Zheng, C., Zhao, Q., 2005. Characteristics and evolution of progressivemetamorphic belt in Chonghuer of the Alati area, Xinjiang. Geoscience 19,334–339 (in Chinese with English abstract).

Yakubchuk, A., 2004. Architecture and mineral deposit settings of the Altaid orogeniccollage: a revised model. Journal of Asian Earth Sciences 23, 761–779.

Yuan, C., Sun, M., Xiao, W.J., Li, X.H., Lin, S.F., Xia, X.P., Long, X.P., Cai, K.D., 2006.Paleozoic accretion of Chinese Altai: geochronological constraints from granitoids.Abstract of Western Pacific Geophysics Meeting, CD-ROM, Beijing, China.

Yuan, C., Sun, M., Xiao, W., Li, X., Chen, H., Lin, S., Xia, X., Long, X., 2007. Accretionaryorogenesis of the Chinese Altai: insights from Paleozoic granitoids. ChemicalGeology 242, 22–39.

Zack, T., Kronz, A., Foley, S.F., Rivers, T., 2002. Trace element abundances in rutiles fromeclogites and associated garnet mica schists. Chemical Geology 184, 97–122.

Zhang, J.H., Wang, J.B., Ding, R.F., 2000. Characteristics and U–Pb ages of zircon inmetavolcanics from the Kangbutiebao Formation in the Altay region, Xinjiang.Regional Geology of China 19, 281–287 (in Chiese with English abstract).

Zhang, H., Niu, H., Terada, K., Yu, X., Sato, H., Ito, J., 2003. Zircon SHRIMP U–Pb dating onplagiogranite from Kuerti ophiolite in Altay, North Xinjiang. Chinese ScienceBulletin 48, 2231–2235.

Zhang, H., Niu, H., Sato, H., Yu, X., Shan, Q., Zhang, B., Ito, J., Nagao, T., 2005a. LatePaleozoic adakites and Nb-enriched basalts from northern Xinjiang, northwestChina: evidence for the southward subduction of the Paleo-Asian Oceanic Plate. TheIsland Arc 14, 55–68.

Zhang, Z., Yan, S., Chen, B., Zhou, G., He, Y., Chai, F., He, L., 2005b. Middle Devonianpicrites of the southern margin of Altay Orogenic Belt and implications for thetectonic setting and petrogenesis. Journal of China University of Geosciences 16,95–103.

Zhang, H.X., Shen, X.M., Ma, L., Niu, H.C., Yu, X.Y., 2008. Geochoronology of the Fuyunadakite, north Xinjiang and its constraint to the initiation of the Paleo-Asian Oceansubduction. Acta Petrologica Sinica 24, 1054–1058.

Zhang, Z., Zhao, G., Santosh, M., Wang, J., Xong, X., Shen, K., 2010. Late Cretaceouscharnockite with adakitic affinities from the Gangdese batholith, southeasternTibet: evidence for Neo-Tethyanmid-ocean ridge subduction? Gondwana Research17, 615–631.

Zheng, C., Kato, T., Enami, M., Xu, X., 2007. CHIME monazite ages of metasedimentsfrom the Altai orogen in northwestern China: Devonian and Permian ages ofmetamorphism and their significance. Island Arc 16, 598–604.

Zhuang, Y.X., 1993. Tectonothermal Evolution in Space and Time and OrogenicProcesses of the Altaide. Jilin Scientific Technical Press, Changchun, China. inChinese with English abstract.

Zindler, A., Hart, S.R., 1986. Chemical geodynamics. Annual Review of Earth PlanetarySciences 14, 493–571.

Zonenshain, L.P., Kuzmin, M.I., Natapov, L.M. (Eds.), 1990. Geology of the USSR: a Plate –Tectonic Synthesis: American Geophysical Union, Geodynamic Series, 21.

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