Mesozoic Sub-Continental Lithospheric Thinning under Eastern Asia (Geological Society Special Publication No. 280)

Mesozoic Sub-Continental Lithospheric Thinning Under Eastern Asia

The Geological Society of London Books Editorial Committee Chief Editor

BOB PANKHURST (UK) Society Books Editors

JOHN GREGORY (UK) JIM GRIFFITHS (UK) JOHN HOWE (UK) PHIL LEAT (UK) NICK ROBINS (UK) JONATHAN TURNER (UK) Society Books Advisors

MIKE BROWN (USA) ERIC BUFFETAUT (France)

RETO GIERE´ (Germany) JON GLUYAS (UK) DOUG STEAD (Canada) RANDELL STEPHENSON (The Netherlands)

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It is recommended that reference to all or part of this book should be made in one of the following ways: ZHAI , M.-G., WINDLEY , B. F., KUSKY , T. M. & MENG , Q. R. (eds) 2007. Mesozoic Sub-Continental Lithospheric Thinning Under Eastern Asia. Geological Society, London, Special Publications, 280. CHEN , B., ZHAI , M.-G. & TIAN , W. 2007. Origin of the Mesozoic magmatism in the North China Craton: constraints from petrological and geochemical data. In: ZHAI , M.-G., WINDLEY , B. F., KUSKY , T. M. & MENG , Q. R. (eds) Mesozoic Sub-Continental Lithospheric Thinning Under Eastern Asia. Geological Society, London, Special Publications, 280, 131 –151.

GEOLOGICAL SOCIETY SPECIAL PUBLICATION NO. 280

Mesozoic Sub-Continental Lithospheric Thinning Under Eastern Asia

EDITED BY

M.-G. ZHAI Chinese Academy of Sciences, China

B. F. WINDLEY The University of Leicester, UK

T. M. KUSKY St. Louis University, USA and

Q. R. MENG Chinese Academy of Sciences, China

2007 Published by The Geological Society London

THE GEOLOGICAL SOCIETY The Geological Society of London (GSL) was founded in 1807. It is the oldest national geological society in the world and the largest in Europe. It was incorporated under Royal Charter in 1825 and is Registered Charity 210161. The Society is the UK national learned and professional society for geology with a worldwide Fellowship (FGS) of over 9000. The Society has the power to confer Chartered status on suitably qualified Fellows, and about 2000 of the Fellowship carry the title (CGeol). Chartered Geologists may also obtain the equivalent European title, European Geologist (EurGeol). One fifth of the Society’s fellowship resides outside the UK. To find out more about the Society, log on to www.geolsoc.org.uk. The Geological Society Publishing House (Bath, UK) produces the Society’s international journals and books, and acts as European distributor for selected publications of the American Association of Petroleum Geologists (AAPG), the Indonesian Petroleum Association (IPA), the Geological Society of America (GSA), the Society for Sedimentary Geology (SEPM) and the Geologists’ Association (GA). Joint marketing agreements ensure that GSL Fellows may purchase these societies’ publications at a discount. The Society’s online bookshop (accessible from www.geolsoc.org.uk) offers secure book purchasing with your credit or debit card. To find out about joining the Society and benefiting from substantial discounts on publications of GSL and other societies worldwide, consult www.geolsoc.org.uk, or contact the Fellowship Department at: The Geological Society, Burlington House, Piccadilly, London W1J 0BG: Tel. þ44 (0)20 7434 9944; Fax þ44 (0)20 7439 8975; E-mail: [email protected]. For information about the Society’s meetings, consult Events on www.geolsoc.org.uk. To find out more about the Society’s Corporate Affiliates Scheme, write to [email protected]. Published by The Geological Society from: The Geological Society Publishing House, Unit 7, Brassmill Enterprise Centre, Brassmill Lane, Bath BA1 3JN, UK (Orders: Tel. þ44 (0)1225 445046, Fax þ44 (0)1225 442836) Online bookshop: www.geolsoc.org.uk/bookshop The publishers make no representation, express or implied, with regard to the accuracy of the information contained in this book and cannot accept any legal responsibility for any errors or omissions that may be made. # The Geological Society of London 2007. All rights reserved. No reproduction, copy or transmission of this publication may be made without written permission. No paragraph of this publication may be reproduced, copied or transmitted save with the provisions of the Copyright Licensing Agency, 90 Tottenham Court Road, London W1P 9HE. Users registered with the Copyright Clearance Center, 27 Congress Street, Salem, MA 01970, USA: the item-fee code for this publication is 0305-8719/07/$15.00. British Library Cataloguing in Publication Data A catalogue record for this book is available from the British Library. ISBN: 978-1-86239-225-0 Typeset by Techset Composition Ltd, Salisbury, UK Printed by Cromwell Press, Wiltshire, UK Distributors North America For trade and institutional orders: The Geological Society, c/o AIDC, 82 Winter Sport Lane, Williston, VT 05495, USA Orders: Tel þ 1 800-972-9892 Fax þ 1 802-864-7626 E-mail [email protected] For individual and corporate orders: AAPG Bookstore, PO Box 979, Tulsa, OK 74101-0979, USA Orders: Tel þ 1 918-584-2555 Fax þ 1 918-560-2652 E-mail [email protected] Website http://bookstore.aapg.org India Affiliated East-West Press Private Ltd, Marketing Division, G-1/16 Ansari Road, Darya Ganj, New Delhi 110 002, India Orders: Tel þ91 11 2327-9113/2326-4180 Fax þ91 11 2326-0538 E-mail [email protected]

Preface One of the most important, still unresolved, tectonic problems of today concerns the remarkable, welldocumented thinning of the subcontinental lithospheric mantle beneath the North China Craton in Eastern China; this is the largest surface area of thinned lithosphere known in the world (500 km  1300 km). Compositional data from xenoliths in Ordovician kimberlites show that the lithosphere – asthenosphere boundary was then at 180 km depth, but similar data from xenoliths in Cenozoic kimberlites and alkali basalts indicate that the boundary was at only 80 km depth at that time. It is well accepted that more than 100 km of lithospheric root is missing under this craton, and was most probably removed in the Cretaceous by still-unknown processes such as delamination, convection, compositional change, or some other mechanism. The overlying crust exposed today contains many diagnostic rocks and structures also largely of Cretaceous age (i.e. granitic rocks; intermediate, mafic and ultramafic rocks including adakites and high-magnesian andesites; metamorphic core complexes; extensional rifts and sedimentary basins; plus gold deposits and other hydrothermal deposits). Abundant seismic, gravity and tomography data provide key constraints on present-day crust –mantle structures. Juvenile Cenozoic alkaline volcanoes and lava fields suggest emplacement of new asthenophere after the thinning event. This volume represents the first compilation in book form of relevant data and ideas on this unique thinning event and craton and global-scale processes responsible for thinning the lithosphere. Our aim in this Special Publication of the Geological Society is to present the first major compilation of multidisciplinary papers that contain new data and ideas that help constrain the thinning problem in Eastern China. Each main paper contains new, unpublished data that provide some constraints. The papers deal with all the main crustal and tectonic results of the thinning process, new geophysical data provide additional constraints on crustal and mantle structures, and several key papers produce new models to explain the evolution of the thinning process. Although most papers result from a meeting in Beijing in June 2005, we have added a few more to provide a balanced coverage of the subject. We divide the book into several sections based on disciplines. An introductory paper by Kusky et al. (a) synthesizes the Precambrian to Cenozoic history of the North China Craton and its margins, to provide the geological and tectonic framework for understanding the thinning in the Mesozoic.

The next section of the book focuses on magmatism and geochemistry, starting with a paper by Zhang on the temporal and spatial distribution of Mesozoic mafic magmatism in the North China Craton. This is followed by a paper by Huang et al. on the contributions of mantle-derived magmas to the lower crust of the craton in the Mesozoic, and then Fan et al. present evidence for chemical and isotopic heterogeneity of the subcontinental lithospheric mantle using data from late Mesozoic magmatic rocks. Guo et al. examine the geochemistry of Mesozoic mafic magmas from the Yanshan belt, and relate this to post-collisional (North China and Mongolia colliding with Siberia) extension. Chen et al. present additional petrological and geochemical data on a wide range of Mesozoic magmas that suggest to them that the magmas were intruded in a back-arc extensional environment induced by the subduction of the palaeo-Pacific plate beneath the North China Craton. The Mesozoic magmas range in age from 180 to 120 Ma, with most crystallizing between 135 and 127 Ma. The next section of the book, on regional structure and tectonics, contains five papers. The first, by Lin et al., describes polyphase Mesozoic structures and tectonics from the Liaoning Peninsula in Eastern China, including Early Triassic contractional structures related to the North China–South China Craton collision, followed by late Mesozoic extension, plutonism, and metamorphic core complex formation, showing a reversal from contractional to extensional structures between 130 and 120 Ma. Li et al. also discuss geological relationships on the Liaoning Peninsula, and show that Palaeoproterozoic plutons of the Jiao-Liao massif experienced partial melting at 160 Ma, and were deformed three times, at 195– 193 Ma, 153–145 Ma, and 135–95 Ma. They relate the partial melting and deformation episodes to collisional events in the Sulu orogen in the early Mesozoic, and palaeo-Pacific subduction throughout the Mesozoic. Shao et al. discuss the Mesozoic sedimentary basins and structures of the Yanshan orogen, and relate these to progressive shallowing of magmatic activity in the mantle between 180 and 130 Ma, and to a tectonic regime change from contraction to extension in the Mesozoic. Cope & Graham describe Late Jurassic to Early Cretaceous rifting and extension following middle to late Mesozoic contraction in NE China. They suggest that the Mesozoic volcanism was related to crustal thickening, and that the Cenozoic rift basins of Eastern China are a more likely upper crustal

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response to delamination, suggesting therefore that loss of the crustal root did not occur until the Cenozoic. Miao et al. describe the Phanerozoic evolution of the Inner Mongolia– Daxinganling orogenic belt (part of the Central Asian orogen) and suggest that Triassic closure of the ocean basins to the north of the craton at the same time as collision in the Qingling–Dabie –Su-Lu orogen to the south may have contributed to the thinning of the lithosphere beneath the North China Craton in the Mesozoic times. Basin evolution and thermal histories form the subject of the next section in the volume. Li et al. describe the basin-fill sequences from the eastern part of the North China Craton, noting that most basins record contractional tectonics before the Late Jurassic or Early Cretaceous, followed by younger extension, reflecting a Mesozoic tectonic reversal that initiated in the north and propagated southward. Hu et al. examine the palaeogeothermal response to late Mesozoic lithospheric thinning using vitrinite reflectance. They conclude that geothermal gradients were highest in the Middle Triassic and Cretaceous, and that the thermal lithosphere thinned from c. 135 km in the early Mesozoic to c. 65 km in the late Mesozoic. The transition from a stable thermal regime to an active one took place at c. 110 Ma, along with surface erosion. Two papers examine the geophysical constraints on loss of the lithospheric root beneath Eastern China. Hao et al. use gravity and tomography models to constrain the structure of the eastern margin of the North China Craton, and interpret their data to define the position of the eastern margin of the craton. Chang et al. relate crustal P-wave velocity distributions to orogenic features and gold deposits, finding that most gold is located in high-velocity areas associated with the Yanshan and Taihangshan ranges, and near the Tan-Lu fault zone. They find that high-temperature

material in the lower crust beneath these regions may have provided the metals that form the upper crustal deposits. Most of the gold in China is located in Mesozoic plutons and structures in the eastern part of the North China Craton. In a paper on the fluid evolution of gold in the Jiaodong (Shandong) Peninsula in Eastern China, Fan et al. describe gold-bearing quartz veins that formed at 120 + 10 Ma, and show that the gold formed during the major Mesozoic tectonic transition from contraction to extension. Thus, loss of the lithospheric root may have driven a major gold-mineralizing event in the upper crust. Deng et al. discuss different geodynamic models for the loss of the lithospheric root, including thermal erosion, delamination and foundering. As a conclusion and summary of the papers in this book, Kusky et al. (b) present a final major analysis of the subject by integrating the data and ideas in the papers of the Special Publication with those recently published in the literature. This synthesis provides an up-to-date, comprehensive and innovative review of the tectonic causes of and crustal and lithospheric responses to the thinning of the subcontinental lithospheric mantle under Eastern Asia in the Mesozoic. This summary demonstrates for the first time the key relationships between the thinning event and the formation of the Solonker orogenic belt, the ultrahigh-pressure Qingling–Dabie Shan–Su-Lu orogenic belt, and the cumulative effects of mantle weakening through the cumulative effects of hydration by multiple subduction episodes beneath the craton. MINGGUO ZHAI BRIAN F. WINDLEY TIMOTHY M. KUSKY QINGREN MENG

Contents Preface Introduction: Precambrian –Palaeozoic history and framework KUSKY , T. M., WINDLEY , B. F. & ZHAI , M.-G. Tectonic evolution of the North China Block: from orogen to craton to orogen Magmatism and geochemistry ZHANG , H.-F. Temporal and spatial distribution of Mesozoic mafic magmatism in the North China Craton and implications for secular lithospheric evolution HUANG , F., LI , S.-G. & YANG , W. Contributions of the lower crust to Mesozoic mantle-derived mafic rocks from the North China Craton: implications for lithospheric thinning FAN , W.-M., GUO , F., WANG , Y.-J. & ZHANG , H.-F. Late Mesozoic mafic magmatism from the North China Block: constraints on chemical and isotopic heterogeneity of the subcontinental lithospheric mantle GUO , F., FAN , W.-M., LI , X.-Y. & LI , C.-W. Geochemistry of Mesozoic mafic volcanic rocks from the Yanshan belt in the northern margin of the North China Block: relations with post-collisional lithospheric extension CHEN , B., ZHAI , M.-G. & TIAN , W. Origin of the Mesozoic magmatism in the North China Craton: constraints from petrological and geochemical data Structures and tectonics LIN , W., FAURE , M., MONIE´ , P. & WANG , Q.-C. Polyphase Mesozoic tectonics in the eastern part of the North China Block: insights from the eastern Liaoning Peninsula massif (NE China) LI , S. Z., KUSKY , T. M., ZHAO , G., WU , F., LIU , J.-Z., SUN , M. & WANG , L. Mesozoic tectonics in the Eastern Block of the North China Craton: implications for subduction of the Pacific plate beneath the Eurasian plate SHAO , J., HE , G. & ZHANG , L. Deep crustal structures of the Yanshan intracontinental orogeny: a comparison with pericontinental and intercontinental orogenies COPE , T. D. & GRAHAM , S. A. Upper crustal response to Mesozoic tectonism in western Liaoning, North China, and implications for lithospheric delamination MIAO , L., ZHANG , F., FAN , W.-M. & LIU , D. Phanerozoic evolution of the Inner Mongolia –Daxinganling orogenic belt in North China: constraints from geochronology of ophiolites and associated formations Basin evolution LI , Z., LI , Y., ZHENG , J.-P. & HAN , D. Late Mesozoic tectonic transition of the eastern North China Craton: evidence from basin-fill records HU , S., FU , M., YANG , S., YUAN , Y. & WANG , J. Palaeogeothermal response and record of Late Mesozoic lithospheric thinning in the eastern North China Craton Geophysical constraints HAO , T.-Y., XU , Y., SUH , M., XU , Y., LIU , J.-H., ZHANG , L.-L. & DAI , M.-G. East Marginal Fault of the Yellow Sea: a part of the conjunction zone between Sino-Korea and Yangtze Blocks? CHANG , X., LIU , Y., ZHAI , M.-G. & WANG , Y. Crustal P-wave velocity distributions and metallotectonics around the North China Craton

vii

1

35 55 77

101

131

153 171

189 201 223

239 267

281 293

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Mineralization FAN , H.-R., HU , F.-F., YANG , J.-H. & ZHAI , M.-G. Fluid evolution and large-scale gold metallogeny during Mesozoic tectonic transition in the Jiaodong Peninsula, eastern China

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Models DENG , J.-F., ZHOU , M.-F., FLOWER , M. F. J., SU , S.-G., ZHAI , M.-G., LIU , C., ZHAO , G.-C., ZHAO , X.-G., ZHOU , S. & WU , Z.-W. A mechanism for transforming buoyant North Chinese cratonic lithosphere to a denser equivalent for delamination

317

Concluding Review KUSKY , T. M., WINDLEY , B. F. & ZHAI , M.-G. Lithospheric thinning in eastern Asia; constraints, evolution, and tests of models

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Index

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Tectonic evolution of the North China Block: from orogen to craton to orogen T. M. KUSKY1, B. F. WINDLEY2 & M.-G. ZHAI3 1

Department of Earth and Atmospheric Sciences, St. Louis University, St. Louis, MO 63103, USA (e-mail: [email protected]) 2

Department of Geology, University of Leicester, Leicester LE1 7RH, UK

3

Key Laboratory of Mineral Resources, Institute of Geology and Geophysics, Chinese Academy of Sciences, Beijing 100029, China

Abstract: The North China Craton contains one of the longest, most complex records of magmatism, sedimentation, and deformation on Earth, with deformation spanning the interval from the Early Archaean (3.8 Ga) to the present. The Early to Middle Archaean record preserves remnants of generally gneissic meta-igneous and metasedimentary rock terranes bounded by anastomosing shear zones. The Late Archaean record is marked by a collision between a passive margin sequence developed on an amalgamated Eastern Block, and an oceanic arc– ophiolitic assemblage preserved in the 1600 km long Central Orogenic Belt, an Archaean– Palaeoproterozoic orogen that preserves remnants of oceanic basin(s) that closed between the Eastern and Western Blocks. Foreland basin sediments related to this collision are overlain by 2.4 Ga flood basalts and shallow marine–continental sediments, all strongly deformed and metamorphosed in a 1.85 Ga Himalayan-style collision along the northern margin of the craton. The North China Craton saw relative quiescence until 700 Ma when subduction under the present southern margin formed the Qingling–Dabie Shan–Sulu orogen (700–250 Ma), the northern margin experienced orogenesis during closure of the Solonker Ocean (500–250 Ma), and subduction beneath the palaeo-Pacific margin affected easternmost China (200–100 Ma). Vast amounts of subduction beneath the North China Craton may have hydrated and weakened the subcontinental lithospheric mantle, which detached in the Mesozoic, probably triggered by collisions in the Dabie Shan and along the Solonker suture. This loss of the lithospheric mantle brought young asthenosphere close to the surface beneath the eastern half of the craton, which has been experiencing deformation and magmatism since, and is no longer a craton in the original sense of the word. Six of the 10 deadliest earthquakes in recorded history have occurred in the Eastern Block of the North China Craton, highlighting the importance of understanding decratonization and the orogen–craton–orogen cycle in Earth history.

The Archaean North China (Sino-Korean) Craton (NCC) occupies about 1.7  106 km2 in northeastern China, Inner Mongolia, the Yellow Sea, and North Korea (Bai 1996; Bai & Dai 1996, 1998; Fig. 1). It is bounded by the Central China orogen (including the Qinling –Dabie Shan–Sulu belts) to the SW, and the Inner Monglia –Daxinganling orogenic belt (the Chinese part of the Central Asian Orogenic Belt) on the north (Figs 1 and 2). The western boundary is more complex, where the Qilian Shan and Western Ordos thrust belts obscure any original continuity between the NCC and the Tarim Block. The location of the southeastern margin of the craton is currently under dispute (e.g. Oh & Kusky 2007), with uncertain correlations between the North and South China Cratons and different parts of the Korean Peninsula. The Yanshan belt is an intracontinental orogen that strikes east –west through the northern part of the craton (Davis et al. 1996; Bai & Dai 1998).

The NCC includes several micro-blocks and these micro-blocks amalgamated to form a craton or cratons at or before 2.5 Ga (Geng 1998; Zhang 1998; Kusky et al. 2001, 2004, 2006; Li, J. H. et al. 2002; Kusky & Li 2003; Zhai 2004; Polat et al. 2005a, b, 2006), although others have suggested that the main amalgamation of the blocks did not occur until 1.8 Ga (Wu & Zhang 1998; Zhao et al. 2001a, 2005, 2006; Liu et al. 2004, 2006; Guo et al. 2005; Kro¨ner et al. 2005a, b, 2006; Wan et al. 2006a, b; Zhang et al. 2006). Exposed rock types and their distribution in these micro-blocks vary considerably from block to block. All rocks .2.5 Ga in the blocks, without exception, underwent the 2.5 Ga metamorphism, and were intruded by 2.5–2.45 Ga granitic sills and related bodies. Nd TDM models show that the main crustal formation ages in the NCC are between 2.9 and 2.7 Ga (Chen & Jahn 1998; Wu et al. 2003a, b). Emplacement of mafic dyke swarms at 2.5– 2.45 Ga has also been

From: ZHAI , M.-G., WINDLEY , B. F., KUSKY , T. M. & MENG , Q. R. (eds) Mesozoic Sub-Continental Lithospheric Thinning Under Eastern Asia. Geological Society, London, Special Publications, 280, 1 –34. DOI: 10.1144/SP280.1 0305-8719/07/$15 # The Geological Society of London 2007.

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T. M. KUSKY ET AL.

80°

70°

90°

100°

120°

110°

130°

140°

50°

CAO

CAO

40°

TM

CCO AH O

NCC

SG

30°

O YC CC

20°

Fig. 1. Simplified map of Asia showing the major tectonic elements. NCC, North China Craton; TM, Tarim Block; CAO, Central Asia orogen; SGO, Songpan Ganzi orogen; CCO, Central China orogen; YC, Yangtze Craton; CC, Cathaysia Craton; AHO, Alpine– Himalaya orogen. Each province has many subdivisions, as discussed in the text.

Fig. 2. Simplified geological map of the North China Craton (after Kusky & Li 2003).

OROGEN CRATON OROGEN CYCLE, NORTH CHINA

3

1.85 Ga Collision of Arc with North China Craton Northern Hebei r Mongolia

Palaeoproterozoic Orogen

Inne

**

** Zunhua * *** Beijing ** Taihang EASTERN Mountain BLOCK

Hengshan Plateau Yinchuan

Wutai Mountain

boundary of North China Craton

2.50 Ga ophiolitic complexes

WESTERN Xian BLOCK Central Orogenic Belt

N 0

km 200

Shenyang West Liaoning Dongwanzi

Kaifeng

* Bangbo

Qinglong foreland basin

2.50-1.8 Ga highpressure granulites 2.50 Ga foreland basin sequences 1.8 Ga granulite: uplifted plateau

Fig. 3. Tectonic map of the North China Craton (modified after Kusky & Li 2003).

recognized throughout the NCC (Liu 1989; Li, J. H. et al. 1996; Li, T. S. 1999). The craton consists of two major blocks (named the Eastern and Western Blocks), separated by the Central Orogenic Belt (Fig. 3). Other blocks, for example the Jiaoliao Block and Alashan Block, have been described (Geng 1998; Zhai 2004), and most appear to have been amalgamated by the time that the Eastern and Western Blocks collided at 2.5 Ga. Some of the boundaries, however, have been reactivated. Wu et al. (1998) suggested that a compositional polarity and diachronous intrusion history in the Eastern Block occurred because an ancient ocean basin between the blocks that now make up the Eastern Block was subducted eastward, beneath the continental block, forming an island arc, which evolved into an arc –continent collisional zone from Honghtoushan, via Qinhuangdao to eastern Shandong. The boundary between the Alashan Block and Western Block is the Western Ordos border fault, the nature of which is not clear. The Western Block (also referred to as the Ordos Block) is a stable part of the craton that has a thick mantle root (based on depth to the low-velocity zone), low heat flow, and has experienced little internal deformation since the Precambrian (Yuan 1996; Zhai & Liu 2003). In contrast, the Eastern Block is unusual for a craton in that it is at

present the site of numerous earthquakes, high heat flow, and a thin lithosphere reflecting the lack of a thick mantle root (Yuan 1996). The NCC is thus one of the world’s most unusual cratons. At one time, it had a typical thick mantle root developed in the Archaean, locally modified at 1.8 Ga, and that was present through the midPalaeozoic as recorded by Archaean-aged mantle xenoliths carried in Ordovician kimberlites (Menzies et al. 1993; Griffin et al. 1998, 2003; Gao et al. 2002; Wu et al. 2003a, b). However, the eastern half of the root appears to have been removed during Mesozoic tectonism. Below we outline the geology of the NCC and surrounding regions, starting with the amalgamation of the craton in the Archaean and/or Palaeoproterozoic and finishing with a summary of the evidence for the distinct behaviour of the Western and Eastern Blocks during Phanerozoic tectonism.

Precambrian geology Major divisions and characteristics of blocks The North China Craton includes a large area of locally well-exposed Archaean crust (Fig. 2),

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including c. 3.8–2.5 Ga gneiss, tonalite –trondhjemite –granodiorite (TTG), granite, migmatite, amphibolite, ultramafic bodies, mica schist, dolomitic marble, graphite- and sillimanite-bearing gneiss (khondalite), banded iron formation (BIF), and meta-arkose (Jahn & Zhang 1984a, b; Jahn et al. 1987; He et al. 1991, 1992; Bai et al. 1992; Bai 1996; Wang 1991; Wang & Zhang 1995; Wang et al. 1997; Wu et al. 1998). The Archaean rocks are overlain by quartzites, sandstones, conglomerates, shales, and carbonates of the 1.85 –1.40 Ga Mesoproterozoic Changcheng (Great Wall) Series (Li et al. 2000a, b). In some areas of the central part of the NCC, 2.40–1.90 Ga Palaeoproterozoic sequences that were deposited in cratonic graben are preserved (Kusky & Li 2003). The North China Craton is divided into two major blocks (Fig. 3) but the boundaries and ages of the intervening orogen have been the subject of some recent debate. One group (e.g. Kusky & Li 2003; Polat et al. 2006) has suggested that the boundary is a Late Archaean–Palaeoproterozoic orogen called the Central Orogenic Belt (COB), that underwent later deformation at c. 1.85 Ga. Other workers (e.g. Zhao et al. 2001a, 2006; Kro¨ner et al. 2006) have suggested that the orogen is a c. 1.85 Ga feature called the Trans North China Orogen (TNCO) that represents collision of the two blocks at 1.85 Ga, and have defined the boundaries as Mesozoic faults. We believe that geological relationships, described below, favour the first division, which is followed here. However, most metamorphic ages demonstrate that strong metamorphism occurred at c. 1.85–1.8 Ga. The Eastern and Western Blocks are separated by the Late Archaean Central Orogenic Belt, in which virtually all U –Pb zircon ages (upper intercepts) fall between 2.55 and 2.50 Ga (Zhang 1989; Zhai et al. 1995; Kro¨ner et al. 1998, 2002; Wilde et al. 1998; Zhao et al. 1998, 1999a, b, 2000, 2001a, b, 2005; Li et al. 2000b; Kusky et al. 2001, 2004; Zhao 2001; Kusky & Li 2003; Polat et al. 2005a, b, 2006). The stable Western Block, also known as the Ordos Block (Bai & Dai 1998; Li et al. 1998), is a stable craton with a thick mantle root, no earthquakes, low heat flow, and a lack of internal deformation since the Precambrian. It has a thick platform sedimentary cover intruded by a narrow belt of 2.55–2.50 Ga arc plutons along its eastern margin (Zhang et al. 1998). Much of the Archaean geology of the Western Block is poorly exposed because of thick Proterozoic and Palaeozoic to Cretaceous platformal cover. A platformal cover on an Archaean basement is typical of many Archaean cratons worldwide. In contrast, the Eastern Block is atypical for a craton in that it has been tectonically active and

has numerous earthquakes, high heat flow, and a thin lithosphere reflecting the lack of a thick mantle root. The Eastern Block contains a variety of c. 3.80 –2.50 Ga gneissic rocks and greenstone belts locally overlain by 2.60 –2.50 Ga sandstone and carbonate units (e.g. Bai & Dai 1996, 1998). Deformation is complex, polyphase, and indicates the complex collisional, rifting, and underplating history of this block from the Early Archaean to the Meso-Proterozoic (Zhai et al. 1992, 2002; Li et al. 2000a; Kusky et al. 2001, 2004; Kusky & Li 2003; Zhai 2004, 2005; Polat et al. 2005a, b, 2006), and again in the Mesozoic –Cenozoic (as described in the papers in this volume). The Central Orogenic Belt includes belts of TTG, granite, and supracrustal sequences that were variably metamorphosed from greenschist to granulite facies. It can be traced for about 1600 km from west Liaoning in the north to west Henan Province in the south (Fig. 3). It should be noted that the COB differs from the TNCO defined by Zhao et al. (2001a). The COB is an Archaean orogen, with Archaean structures defining its boundaries, whereas the TNCO is defined as a Proterozoic orogen, albeit one bound by Mesozoic structures. High-grade regional metamorphism, including migmatization, occurred throughout much of the Central Orogenic Belt between 2.60 and 2.50 Ga (Zhai 2004), with final uplift of the metamorphic belt during c. 1.90 – 1.80 Ga extensional tectonism (Li et al. 2000a) or a collision on the northern margin of the NCC (Kusky & Li 2003). Greenschist- to amphibolitegrade metamorphism predominates in the southeastern part of the COB (such as in the Qinglong belt, Fig. 2), but the northwestern part is dominated by amphibolite- to granulite-facies rocks, including some high-pressure assemblages (10– 13 kbar at 850 + 50 8C; Li et al. 2000b; Zhao et al. 2001a, b; see additional references given by Kro¨ner et al. 2002). The high-pressure assemblages occur in the linear Hengshan belt (Fig. 4), which extends for more than 700 km with a ENE – WSW trend. Internal (western) parts of the orogen are characterized by thrust-related subhorizontal foliations, shallow-dipping shear zones, recumbent folds, and tectonically interleaved highpressure granulite migmatite and metasedimentary rocks. The COB is in many places overlain by sedimentary rocks deposited in graben and continental shelf environments, and is intruded by c. 2.5–2.4 and 1.9 –1.8 Ga dyke swarms. Several large 2.2–2.0 Ga anorogenic granites have also been identified within the belt (Li & Kusky 2007). Recently, two linear zones of deformation have been documented within the belt, including a high-pressure granulite belt in the west (Li et al.

OROGEN CRATON OROGEN CYCLE, NORTH CHINA

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Fig. 4. Simplified geological map of the Hengshan– Wutaishan– Fuping area, showing relationships between high-pressure granulites and gneiss north of the shear zone on the north side of Hengshan Mountains, medium-pressure granulites to the south, and amphibolite- to greenschist-facies rocks of the Wutai Group and greenstone belt. Map modified after Yuan (1988) and Li et al. (2004).

2000a), and a foreland basin and fold– thrust belt in the east (Li, J. H. et al. 2002; Kusky & Li 2003; Li & Kusky 2006). The high-pressure granulite belt is separated by normal faults from the Western Block,

which is overlain by thick metasedimentary rocks (khondalites) that are younger than 2.40 Ga, and were metamorphosed at 1.862.7 + 0.4 Ga; A. Kro¨ner, pers. commun.).

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High-pressure granulites The Hengshan high-pressure granulite (HPG) belt consists of several metamorphic terranes, including the Hengshan, Huaian, Chengde, West Liaoning, and Southern Taihangshan metamorphic complexes (Figs 2–4). The HPG commonly occurs as isolated pendants within intensely sheared TTG (2.60– 2.50 Ga) and granitic gneiss (2.50 Ga), and is widely intruded by 2.20–1.90 Ga K-granite and mafic dyke swarms (2.45–2.40 Ga, 1.77 Ga) (Li et al. 2000b; Kro¨ner et al. 2002; Peng et al. 2007). Locally, thrust slices of lower metamorphic grade khondalite and metamorphosed turbiditic sediments are interleaved with the high-pressure granulite rocks. The main rock type of the complexes is a garnet-bearing mafic granulite with characteristic plagioclase –orthopyroxene coronas surrounding the garnets, which show evidence for rapid exhumation-related decompression (at c. 1.9–1.8 Ga) from peak P–T of 1.2–0.9 GPa and 700– 800 8C (Zhao et al. 2000; Kro¨ner et al. 2002). At least three types of REE patterns are shown by the mafic rocks from flat to moderately light REE (LREE)-enriched, indicating original crystallization in a continental margin or island-arc setting (Li, J. H. et al. 2002). The subsequent highpressure metamorphism occurred during pre-2.5 Ga partial subduction of the mafic rocks, which was then followed by collision and the rapid rebound– extension that is recorded by 2.50–2.40 Ga mafic dyke swarms and graben-related sedimentary rock sequences in the Wutai Mountains–Taihang Mountains areas (Kusky & Li 2003; Kusky et al. 2006). Another kind of high-pressure granulites occur as deformed and pulled-apart dykes. They yield sensitive high-resolution ion microprobe (SHRIMP) zircon ages of 1973 + 4 Ma and 1834 + 5 Ma, with a core residual age of 2.0–2.1 Ga (Peng et al. 2005, 2007). Zhao et al. (2001a, b, 2005, 2006), Wilde et al. (2003), and Kro¨ner et al. (2005a, b, 2006) have suggested that the c. 1.9–1.8 Ga granulite event in the NCC is related to the continent–continent collision between the Eastern and Western Blocks of the craton. This model is supported by the interpretation of clockwise metamorphic P–T –t paths that show crustal thickening related metamorphism at 1.85 Ga, in support of a collision at this time. However, Kusky & Li (2003) noted that the structural, sedimentological, and geological field data suggested collision of the Eastern and Western Blocks at 2.5 Ga, and that the 1.9–1.8 Ga granulite event occurs throughout rocks across the entire northern half of the craton, not just in the COB, and that it might be related to a collision along the northern margin of the craton, forming

an east –west orogen by 1.8 Ga. O’Brien et al. (2005) recognized two main types of granulites, including high-pressure mafic granulites in the north, and medium-pressure granulites in the south, separated by the east –west-striking Zhujiafang shear zone. Further south, metamorphic facies are even lower grade, dominated by amphibolite to greenschist facies in the Wutaishan (O’Brien et al. 2005), providing evidence for north to south crustal staking of higher over lower grade rocks at c. 1.9–1.8 Ga. Santosh et al. (2006) have related ultrahigh-temperature metamorphism (975 8C at 9 kbar, and 900 8C at 12 kbar) at 1927 + 11 Ma, and 1.1819 + 11 Ma, to the formation of a 1.9–1.8 Ga collisional orogen along the north margin of the NCC during the amalgamation of the Columbia supercontinent.

2.5 Ga foreland basin The Late Archaean Qinglong foreland basin and fold–thrust belt (Fig. 3) trends north– south to NE– SW, and is now preserved as several relict folded sequences (Kusky & Li 2003; Li & Kusky 2006). Its general sedimentary rock sequence from bottom to top can be further divided into three subgroups of quartzite–mudstone–marble, turbidite, and molasse. The lower subgroup, of quartzite – mudstone–marble, is well preserved in central sections of the Qinglong foreland basin (Taihang Mountains), which includes numerous shallowly dipping structures, and is interpreted to be a product of pre-2.5 Ga passive margin sedimentation on the Eastern Block. It is overlain by lower-grade turbidite and molasse-type sediments. The western margin of the Qinglong foreland basin is intensely reworked by thrusting and folding, and is overthrust by rocks of an active margin (TTG gneiss, ophiolite fragments, accretionary wedge type metasediments). To the east, rocks of the basin are less deformed, defining a gradual transition from highgrade metamorphism and ductile structures of the COB to an upper crustal level fold–thrust belt then foreland basin style structures to the east. The passive margin sedimentary rocks and the Qinglong foreland basin are intruded by a c. 2.40 Ga diorite and gabbroic dyke complex (Li & Kusky 2006), and are overlain by graben-related sedimentary rocks and 2.4 Ga flood basalts. In the Wutai and North Taihang basins, many ophiolitic blocks are recognized along the western margin of the foreland fold-and-thrust belt. These typically consists of pillow lava, gabbroic cumulates, and harzburgite, with the largest block being 10 km long in the Wutai –Taihang Mountains (Wang et al. 1997).

OROGEN CRATON OROGEN CYCLE, NORTH CHINA

Timing of collisional orogenesis in the Central Orogenic Belt Whereas it is well recognized that the Central Orogenic Belt records the collision between the Western and Eastern Blocks of the NCC, the timing of this collision is debated. Zhao and co-workers (Zhao et al. 2001a, b, 2005, 2006; Kro¨ner et al. 2006) suggested that collision between the Western and Eastern Blocks of the NCC occurred at 1.8 Ga, based on the metamorphic ages of high-pressure granulites and their inferred isothermal decompression (ITD) type clockwise P –T paths. ITD type P –T paths in regionally metamorphosed rocks are generally interpreted as reflecting double thickening of crust followed by erosion and uplift. Thus, in the Zhao et al. scenario, a continental arc that had been active on the western edge of the Eastern Block since 2.5 Ga was transformed to a continent– continent collision zone at c. 1.85 Ga with the collision of the passive margin of the Western Block, indicating a life span for this margin of 650 Ma. However, many U –Pb and other metamorphic ages point to a major amphibolite– granulite-facies event at 2.5 Ga (Kro¨ner et al. 1998; Zhai & Liu 2003; Kusky et al. 2006), a feature not accounted for in the Zhao et al. model. Several other aspects of the Zhao et al. model make it untenable. First, the proposition of having an active margin for 650 Ma is unlikely, especially when the geological record in the NCC shows little evidence for any accretionary activity in this period. Such a long-lived accretionary margin would be expected to produce an accretionary orogen on the scale of the Makran or the southern Alaska margin, yet the proposed location of the margin preserves no such rocks. Further, in the Zhao et al. (2006) interpretation, the granulites along the northern margin of the craton are explained by the unlikely scenario in which the two continental blocks both independently developed granulitefacies belts on one of their margins, which fortuitously became perfectly lined up to form one continuous belt along the northern margin of the craton at 1.8 Ga. The Zhao et al. model relies on the interpretation of the significance of c. 1.85 Ga metamorphic ages and P–T –t paths from a major event at 1.85 Ga. Recent detailed mapping, analysis of structures, sedimentary basins, and the distribution of tectonic belts or rocks types in the craton suggest that there are other possible interpretations of the 1.85 Ga event. Furthermore, other workers (e.g. Li et al. 1996, 2000a, b; O’Brien et al. 2005; Santosh et al. 2006, and references therein) have shown that the ultra high-temperature and high-pressure granulites are distributed across the

7

northern part of the craton, and not confined to the Central Orogenic Belt. Kusky and coworkers (Kusky et al. 2001, Kusky 2004; Kusky & Li 2003; Polat et al. 2005a, b, 2006) suggested that the Eastern and Western Blocks collided at 2.5 Ga, forming a 200 km wide orogen that included development of a foreland basin on the Eastern Block, and a granulite-facies belt on the Western Block. Evidence for this collision is found as remnants of 2.5 Ga oceanic crust (Kusky et al. 2001; Kusky 2004; Polat et al. 2005a, b, 2006), island arcs, accretionary prisms, and deformed continental fragments, which show a consistent 2.5 Ga metamorphism. Late Archaean collision was, in this scenario, followed by postorogenic extension and rifting that led to the emplacement of mafic dyke swarms and development of extensional basins along the COB, as well as to the opening of a major ocean along the northern margin of the NCC (Kusky & Li 2003).

1.85 Ga continent – continent collision on the northern margin of the craton After collision at c. 2.5 Ga and post-collisional extension by 2.4 Ga, the North China Craton was in a relatively inactive tectonic stage with the exception of deformation, magmatic activity and metamorphism associated with an Andean-type margin that was active on the north margin of the craton from 2.2 to 1.85 Ga. Then an important metamorphic event happened between 1900 and 1800 Ma. As a result, all Precambrian rocks of the craton experienced the same metamorphic episode at 1900–1800 Ma, and associated migmatization and intrusion of crustal melt granites. Kusky & Li (2003) related this event to a continental collision on the northern margin of the craton, associated with the formation of a new east– west-striking foreland basin (in which the Changcheng Series of conglomerates, sandstones and shales was deposited), and was followed closely by a new period of post-orogenic extension. High-pressure granulites were developed in an east–west belt in the north (the Inner Mongolia–Eastern Hebei Palaeoproterozoic orogen), with polyphase granulites preserved from UHT processes in the Andean-type arc, and where the east –west belt crosses the COB. Alternatively, Zhai (2004) proposed that the c. 1.8 Ga event represents a continental geological process within the craton: an upwelling mantle plume caused uplift of the craton basement as a whole and was closely followed by the development of an aulacogen system. A series of continent rifts were developed, with alkalic volcanic eruption and intrusion of anorogenic magmatic association

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(rapakivi– anorthosite–gabbro) and mafic dyke swarms. The Mesoproterozoic sedimentary sequences in the Yanshan rift are called the Changcheng–Jixian System, which was deposited at c. 1800–1500 Ma. However, the age of the upper Jixian System is not defined: it could extend to c. 1400 –1100 Ma. Zhao et al. (2004) suggested that the volcanic eruption centre of the rift system was in western Henan Province. From c. 1800 Ma to 1700 Ma (the Xiong’er Group), the rift extended to the west, east and north, forming a triple junction. Finally, dioritic intrusions indicate rifting-end magmatic activity. The rift system mainly trends NE–SW to east– west and branches off into the Taihang Mountains to the south. The northern margin of the craton remained episodically active as a convergent–accretionary margin (separated by periods of passive margin sedimentation) for the next several hundred million years, growing northward and accommodating the southward(?) subduction of thousands of kilometres of oceanic lithosphere. The 1.8 Ga event that formed the high-pressure granulites with clockwise P–T paths was interpreted by Kusky & Li (2003) as being related to a (continental?) collision outboard of the Inner Mongolia– Eastern Hebei orogen, and closure of a back-arc basin preserved along the north margin of the craton. Following collision at 1.85 Ga, extensional tectonics gave rise to a series of aulacogens and rifts that propagated across the craton, along with the intrusion of mafic dyke swarms. On the northern margin of the craton at Bayan Obo, a basement of migmatites is overlain unconformably by a 2 km thick shelf sequence of c. 2.07 –1.5 Ga quartzites, shales, limestones, dolomites and conglomerates. Carbonatite dykes (Le Bas et al. 1992; Fan et al. 2002) emplaced into the sedimentary rocks are associated with the largest REE deposit in the world that has a Sm– Nd mineral age of 1426 + 40 Ma and a monazite age of 1350 + 149 Ma (Nakai et al. 1989). On the southwestern margin of the NCC the Western Block gneisses and migmatites are overlain by marbles and intruded by the Jinchuan lherzolite body, which contains the third largest nickel deposit in the world (Chai & Naldrett 1992). Troctolite associated with the lherzolite has a 206Pb/238U SHRIMP age on zircons of 827 + 8 Ma, regarded by Li et al. (2004) as the crystallization age of the ultramafic intrusion. As Li et al. suggested, the Jinchuan intrusion may have been emplaced as a result of mantle plume activity during the break-up of the Rodinia supercontinent. Many relationships between Palaeoproterozoic volcanosedimentary groups and basement blocks in the eastern part or the craton are still enigmatic. For instance, rocks of the Liaohe Group on the

Jiadong Peninsula, and the Guanghua, Ji’an and Liaoling groups in Jilin Province, have been assigned various ages ranging from 2.5 to 1.9 Ga, and their tectonic environments have been interpreted as accretionary prism, collision-related, and rift related (e.g. see Zhai 2005; Li et al. 2006; Lu et al. 2006). Very little structural work has been published on these rocks, and it is clearly needed to understand the role of these rock groups in the tectonic evolution of the craton. From the late Neoproterozoic until the end of the Palaeozoic, the NCC behaved as a coherent, stable continental block, as evidenced by deposition of shallow-marine carbonate platform sediments throughout the Palaeozoic (e.g. Metcalfe 1996, 2006). Breaks in sedimentation, however, were associated with deformation and orogeny along all margins of the craton and a regional disconformity between the Upper Ordovician and Upper Carboniferous units (Wang 1985). The latter may have resulted from the global eustatic lowstand of sea level following the early Palaeozoic orogeny or from double-vergent subduction beneath the north and south margins of the craton (the Qaidam plate was subducted beneath the southern margin of the craton, and several oceanic plates subducted beneath the north margin of the craton (Yin & Nie 1996)). Moreover, it is during this interval that diamond-bearing kimberlites erupted in several areas of the Eastern Block of the NCC (Fig. 2; Menzies et al. 1993; Griffin et al. 1998). The diamonds and the P –T array inferred from garnets carried in these kimberlites testify to the presence of a thick (170 km) lithospheric keel, similar to that observed in Archaean cratons elsewhere (e.g. Kaapvaal, Slave, Siberia; see Menzies et al. 1993; Griffin et al. 1998, 2003).

Phanerozoic tectonics Major orogenic belts, faults and basins It is fair to say that the detailed geological and tectonic histories of the margins of the NCC are, for the most part, very poorly understood. Using current palaeomagnetic data, de Jong et al. (2006) suggested that in the Early Palaeozoic the NCC, South China (Yangtze) Craton and the Tarim Craton (Fig. 1) were microcontinents fringed by subduction–accretion complexes and island arcs along the northeastern Cimmerian margin of Gondwana (Fig. 5). Rifting in the Early Carboniferous was followed by drifting of the Precambrian blocks across the Palaeo-Tethys Ocean, and their amalgamation to form much of what is now China in Permo-Triassic times. The Solonker and Dabie sutures (see Figs 2, 6 and 7) record respectively

OROGEN CRATON OROGEN CYCLE, NORTH CHINA

Fig. 5. Palinspastic map and schematic cross-sections showing the evolution of the North China Craton in the Palaeozoic. Modified after Heubeck (2001) and Yue et al. (2001). AT, Altyn Tagh; BA, Baoerhantu arc; DA, Dongqiyishan arc; DUA, Don Ujimqin arc; HGS, Hegenshan suture; HM, Hanshan microcontinent; HS, Hongshishan suture; MSQ, middle and south Qilian; NAS, North Altyn Tagh suture; NC, North China Craton; NETB, northeastern Tarim Block; NQS, north Qilian suture; SLS, Solon–Linxi suture; XM, Xilin Hot microcontinent; XS, Xiaohuangshan suture; YA, Yuanbaoshan arc. It should be noted that although the NCC and Tarim Block experienced craton margin tectonism throughout the Palaeozoic, the craton interior was relatively quiescent. However, subduction of thousands of kilometres of oceanic lithosphere under the craton from the Palaeotethys in the south, and Turkestan (Palaeoasian) Ocean strands in the north, significantly hydrated and weakened the subcontinental lithospheric mantle, perhaps creating conditions favourable for root loss in the Mesozoic.

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Fig. 6. Schematic map of the northern margin of the North China Craton, including the Inner Mongolia–Northern Hebei Palaeoproterozoic orogen, and the Central Asian orogen (modified after Xiao et al. 2003). The Solonker suture marks the composite suture between terranes accreted to the northern margin of the North China Craton, and terranes accreted to the southern margin of the Siberian Craton.

10 T. M. KUSKY ET AL.

Fig. 7. Map of the Qingling–Dabie orogen (after Li, S. Z. et al. 2006). The two sutures in the orogen, including Shangdan suture in the north, and the Mianlue suture in the south, should be noted. The Shangdan suture resulted from Middle Palaeozoic closure of the Shangdan ocean and collision of the North China Craton and the Qinling– Dabie microplate. The Mianlue suture, however, resulted from Late Triassic closure of the Mianlue ocean and collision of the Qinling –Dabie microplate and the South China Craton. Map drawn by S. Z. Li. Abbreviations in inset map are as in Figure 1.

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terrane accretion from the north (during closure of the Turkestan Ocean) and collision of the South China Craton with the NCC in the south (e.g. Li et al. 1995; Metcalfe 1996). The main Mesozoic events to affect the NCC are traditionally referred to as the Late Triassic –Early Jurassic Indosinian orogeny, and the Late Jurassic–Early Cretaceous Yanshanian orogeny (Yang et al. 1986). Main surface features related to these events include major east –west and north –south fold belts, widespread plutonism, and extensional faults. The structural history of the relatively flat-lying Palaeozoic sedimentary cover of the NCC shows that it was stable until Jurassic times (Wang 1985) although deformation on the craton margins began earlier. Kimberlites found in the Taihang – Luliang regions are Mesozoic –Tertiary in age and are related to uplift of the Shanxi highlands in the centre of the craton, which preceded and represents early stages of the young rifting in this area (Ke & Tian 1991; Dobbs et al. 1994; Zheng et al. 1998, 2001). On the eastern side of the craton, one of the world’s largest continental margin transcurrent faults, the Tan-Lu fault, constitutes the most striking structural feature of the region (Fig. 2). It stretches more than 1000 km subparallel to the Pacific margin and probably extends into Russia (Xu & Zhu 1994). The timing of early motion and cause of formation of the Tan-Lu fault are controversial. Various workers have proposed Triassic (Okay & Sengo¨r 1992; Yin & Nie 1993) or Cretaceous (Xu et al. 1987; Xu 1993; Xu & Zhu 1994) ages for initial motion, reflecting initiation either from collision between South China (Yangtze) Cratons and the NCC (Okay & Sengo¨r 1992; Yin & Nie 1993) or from oblique convergence between the Pacific and Asian plates (Xu et al. 1987). The apparent offset of the Dabie Shan and Su-Lu ultrahigh-pressure rocks suggests c. 500 km of initial sinistral motion on the Tan-Lu fault during the Triassic –Jurassic collision of the North and South China Cratons (Okay & Sengo¨r 1992; Yin & Nie 1993). However, the central part of the fault indicates c. 740 km of sinistral displacement (Xu et al. 1987). Large-scale left-lateral strike-slip motion occurred on the Tan-Lu fault at c. 132 – 128 Ma (Early Cretaceous). Geological evidence of Early Jurassic to midCretaceous tectonism in the NCC is abundant, and not just recorded along the Tan-Lu fault system. Widespread 147– 112 Ma magmatism included the intrusion of adakites, reflecting subduction of perhaps as many as three distinct slabs (Xu 1990; Zhang, L. C., et al. 2000; Davis et al. 2001; Wang et al. 2001; Zhang, Q., et al. 2001; Wei et al. 2002; Xu et al. 2002; Davis, 2003; see also Castillo 2006). The formation of China’s most important

gold vein deposits occurred at the same time along the northern, eastern, and southern margins of the Eastern Block (Mao et al. 1999; Zhou et al. 2002; Yang, J.-H. et al. 2003; Fan et al. 2007). Unroofing of many metamorphic core complexes (c. 140–105 Ma), products of SE–NE extension (Niu 2005; Zhang et al. 1994; Zheng et al. 1998, 2001; Zhang, Y. Q. et al. 1998; Webb et al. 1999; Zhang, Q., et al. 2001; Davis et al. 2002; Darby et al. 2006; Li et al. 2007), and major animal extinctions were also significant in this period (Chen et al. 1997; Wang et al. 2001). These observations support a change from a relatively internally stable craton, from c. 1900 to 250 Ma, to a middle to late Mesozoic situation where the margins of the Eastern Block underwent significant Yanshanian orogenesis. This tectonism reflects three relatively contemporaneous collisional or subduction events, or both: (1) the collision of the Yangtze Craton to the south; (2) the closure of the Turkestan Ocean (forming the Solonker suture) and accretion of the oceanic arcs on the north; (3) and oblique subduction of Palaeopacific oceanic crust on the east (Fig. 5). Below, we discuss each of these settings.

Northern margin: the Solonker suture, and Palaeozoic subduction beneath the north margin of the NCC The Palaeoasian or Turkestan Ocean was present on the northern side of the NCC throughout the Palaeozoic, with Palaeo-Tethys to the south (e.g. Metcalfe 1996, 2006). Several subduction zones were active during this interval, leading to continental growth through accretion of terranes along the northern margin of the craton and the generation of arc magmas (Davis et al. 1996, 2002, 2006; Yue et al. 2001; Xiao et al. 2003). These terranes north of the NCC (Fig. 6) host more than 900 Late Palaeozoic to Early Triassic plutons (Sengo¨r et al. 1993; Sengor & Natal’in 1996; Xiao et al. 2003). Xiao et al. (2003) suggested that these plutons are related to closure of the Palaeoasian ocean at the end of the Permian. Closure is marked by the Solonker suture (Fig. 6) and 300–250 Ma south-directed subduction beneath the accreted terranes along the northern side and the northern margin of the NCC itself (Xiao et al. 2003). Continued convergence from the north during Triassic and Jurassic times caused post-collisional thrusting and considerable crustal thickening on the NW side of the craton (Xiao et al. 2003). The northeastern margin of the NCC with a Permian shelf sequence collided with the Khanka Block in Late Permian to Early Triassic times, as indicated by syncollisional granites (Jia et al. 2004). Many of the subsequent later Mesozoic

OROGEN CRATON OROGEN CYCLE, NORTH CHINA

granitoids, metamorphic core complexes, and extensional basins, south of the Solonker suture in the northern part of the NCC and the adjacent Palaeozoic accretionary orogen (Fig. 6), may be related to post-collisional Jurassic – Cretaceous collapse of the massive Himalayan-style Solonker orogen and plateau (Ritts et al. 2001; Xiao et al. 2003; Gregory et al. 2006).

The south: Qingling –Dabie Shan– Sulu orogen The Qinling –Dabie orogen is marked by the terranes forming the irregular suture between the NCC and South China Craton (Fig. 7). It is a major part of the east –west-trending Central China orogen (Jiang et al. 2001), which extends for 1500 km eastward from the Kunlun Range to the Qinling Range, and then 600 km farther east through the Tongbai– Dabie Range. Its easternmost extent, offset by movement along the Tan-Lu fault system, continues northeastward through the Sulu area of the Shandong Peninsula and then into South Korea. Ratschbacher et al. (2003) suggested that the Sulu belt continues through the Imjingang fold belt of Korea, yet the presence of 230 Ma eclogites in the southern Gyeonggi massif (Oh 2006; Oh & Kusky 2007) suggests that the Sulu belt may alternatively extend through South Korea. The intermittent presence of ultrahigh-pressure diamonds, eclogites and felsic gneisses indicates very deep subduction along a cumulative .4000 km long zone of collisional orogenesis (Yang, J. S., et al. 2003). The rifting and collisional history throughout the Palaeozoic of the NCC with blocks and orogens to the south, such as the North Qinling terrane, the South Qinling terrane, and eventually (in the Triassic) the South China Precambrian block, is complicated and controversial (Meng & Zhang 1999). In the Early Palaeozoic, northward subduction of the Qaidam –South Tarim plate (possibly connected with the South China plate) took place beneath the active southern margin of the NCC (Li, S. Z. et al. 2002, 2006b). The NCC, probably together with the Tarim Block, collided with the South Tarim –Qaidam Block in the Devonian, then with the South China Block in the Permo-Triassic (Li, S. Z. et al. 2006b, and references therein). This latter collision resulted in exposure of ultrahighpressure rocks from c. 100 km depth in Dabie Shan, and westward escape of the South Tarim– Qaidam Block (e.g. Sengo¨r 1985; Yang et al. 1986; Yin & Nie 1996; Hacker et al. 2000; Ratschbacher et al. 2000, 2003), and caused uplift of the large Huabei plateau in the eastern NCC (Fig. 7). Younger extrusion tectonics related to Himalayan

13

collisions further west resulted in c. 500 km of left-lateral motion along the Altyn– Tagh fault, separating the NCC from the South Tarim –Qaidam Block, slicing and sliding to the west the arc that formed on the southern margin of the NCC during Early Palaeozoic subduction (Fig. 8). The terrane accretion and eventual continent– continent collision along the southern margin of the NCC are defined by a geometrically irregular suture, defining a diachronous convergence with a complex spatial and temporal pattern (e.g. Tapponnier et al. 1982; Yin & Nie 1993; Li, S. Z. et al. 2006b). Many models of extrusion tectonics, such as eastward, vertical (upward), and lateral, have been proposed in the last decade for the Qinling –Dabie orogen (Hacker et al. 2000; Li, S. Z. et al. 2002; Wang et al. 2003). Maruyama et al. (1994) proposed that vertical extrusion was important to Triassic exhumation of the ultrahigh-pressure rocks in the eastern part of the orogen. Hacker et al. (2000) pointed out that an orogen-parallel, eastward extrusion occurred diachronously between 240 and 225–210 Ma. Ratschbacher et al. (2000) described Cretaceous to Cenozoic unroofing that was initially dominated by eastward tectonic escape and Early Cretaceous Pacific back-arc extension, and then mid-Cretaceous Pacific subduction. Wang et al. (2003) proposed that the Triassic Dabie highpressure –ultrahigh-pressure metamorphic rocks were originally beneath the Foping dome, which is in the narrowest part of the Qinling Belt, and that these rocks were extruded eastward to their presentday location. We also suggest that the root loss event beneath the adjacent NCC was related to the continental- scale tectonism in the Dabie –Qingling orogen. It is probably more than a coincidence that two of the most unusual tectonic events in the geological record (root loss under the NCC and ultrahigh-pressure metamorphism in Dabie Shan) are geographically and temporally coincident.

The east: Pacific plate subduction Subduction along the Pacific margin of the NCC (Fig. 8) was active from 200 to 100 Ma, starting soon after closure of the ocean basins on the northern side of the craton (Heubeck, 2001; Xiao et al. 2003). Westward-directed oblique subduction was responsible for the generation of arc magmas, deformation, and possibly mantle hydration during this interval (Xu 1990). Although the duration and history of Mesozoic subduction beneath the eastern margin of the NCC is not well known, the active margin stepped outwards by Cenozoic times (Fig. 9), from when a better record is preserved. Numerous plate reconstructions (e.g. Engebretson et al. 1985; Stock & Molnar 1988; Hall 1997) for the Cenozoic of Asia and the

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90°

100° 110° 120° 130°

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SIBERIA

al

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70° subduction zone reverse fault strike-slip fault normal fault extension compression block motion

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IN AW A

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20° PHILIPPINE SEA PLATE

INDIA SOUTH CHINA SEA

10° 500 km 10° 80°

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Fig. 8. Tectonic map of Asia (modified after Zhang, Y. Q. et al. 2003a), showing relationships between the India–Asia collision, escape of Indonesian and South China blocks seaward, and extension from Siberia to the Pacific margin. (Note also the opening of back-arc basins including the Sea of Japan and the South China Sea, and extension in the Bohai Basin and eastern part of the NCC.) The North China Craton is also strongly influenced by Pacific and palaeo-Pacific subduction, perhaps also inducing extension in the eastern NCC. The palaeo-Pacific and Pacific subduction zones developed in the Mesozoic, and also contributed to the hydration of the subcontinental lithospheric mantle beneath the NCC.

Eastern Pacific basin (Fig. 9) show that a wide scenario of different plates, convergence rates, and angles of subduction definitely relate to some of the processes of basin formation, magmatism, and deformation in the easternmost NCC (e.g. Northrup et al. 1995; Hall 1997; Li 2000; Li et al. 2007). The implication of long-lived subduction beneath the NCC is important. When oceanic lithosphere subducts, it dehydrates and thereby weakens the upper mantle. It lowers the melting temperature (solidus), and decreases the mantle viscosity. Only 100– 1000 ppm additional water decreases mantle viscosity by two orders of magnitude (Niu 2005; Komiya & Maruyama 2006). According to Komiya & Maruyama (2006) this is the principal

cause of the fragmentation of the oceanic lithosphere in the Western Pacific. The idea that subduction of water into the mantle caused hydroweakening of the subcontinental lithosphere and was responsible for the thinning–delamination under the Eastern Block of the North China Craton came independently from Niu (2005) and Windley et al. (2005). However, whereas Niu (2005) considered that subduction by the Pacific plate was sufficient to carry water to the upper mantle, Windley et al. (2005), building on the ideas of Maruyama et al. (2004) and Komiya & Maruyama (2006) of double subduction, as summarized above, extended the process to include subduction zones sited on the Solonker, Dabie Shan and Mongol –Okhotsk sutures.

OROGEN CRATON OROGEN CYCLE, NORTH CHINA

15

Fig. 9. Palinspastic maps showing the possible plate interactions along the Pacific margin of the NCC in the Mesozoic. (Note active subduction and episodes of ridge subduction).

Liu et al. (2001) established a connection between volcanic activity and extension in NE and Eastern China from c. 86 Ma to the present and the younger opening of the Japan Sea. However, the area of delamination under the Eastern Block of the NCC was also subjected to earlier subduction from the Solonker Ocean to the north and Dabie Ocean to the south, as described above, and the Cenozoic northerly subduction of the Indo-Australian plate. It is thus difficult to specifically target one major subduction event as the cause of many of the major deformational features. In fact, more different oceanic lithosphere fragments have probably been subducted under

the eastern NCC than under any other Phanerozoic continental block, which may have extensively hydro-weakened the upper mantle (e.g. Niu 2005). Windley et al. (2005) suggested that Jurassic orogenic collapse at the northern and southern margins of the craton triggered the delamination. In a similar model, Zhang et al. (2003) proposed that Palaeozoic subduction of ocean crust beneath both the northern and southern margins of the NCC was responsible for destabilization of the eastern NCC and the resulting thinning and replacement of the lithospheric mantle. However, they envisaged the northern subduction zone as being sited on the margin of the Mongol–Okhotsk

16

T. M. KUSKY ET AL.

Ocean, which would be hundreds of kilometres north of the Solonker Ocean and the preferred site in the present study.

From contraction to extension The tectonics of much of Asia changed from contractional to extensional at c. 130– 120 Ma, and this could be the best approximation for the time of the original subcontinental mantle root loss beneath the NCC. Meng (2003) and Meng et al. (2003) suggested that the Jurassic collision of the amalgamated North China–Mongolia Block with the Siberian plate (Fig. 6) that gave rise to the Mongol–Okhotsk suture led to formation of a high-standing plateau. Gravitational collapse of the thickened crust led to Late Jurassic–Early Cretaceous crustal extension throughout the orogenic belts of Southern Mongolia and Northern China, and coeval thrusting to form the Yanshan belt on the northern margin of the NCC (e.g. Davis et al. 1996). This model, however, ignores the more southerly Solonker suture and associated Late Permian closure of the Palaeoasian Ocean near the Mongolia–China border. This Siberia–Mongolia collision with the simultaneously amalgamating Chinese Precambrian blocks gave rise to a major Himalayan-style orogen or even plateau, the post-collisional collapse of which was probably responsible for the Jurassic thrusting and for the formation of Cretaceous basins and metamorphic core complexes (Xiao et al. 2003). The Late Jurassic Yanshanian orogen (Fig. 10) formed in response to the closure of the Palaeoasian Ocean along the north margin of the NCC, subduction of the palaeo-Pacific plate beneath the eastern margin of the NCC, and continued convergence between the NCC and South China Block in the south. This three-sided convergence in the Late Jurassic during the Yanshanian orogeny resulted in further uplift of the Huabei plateau, and widespread deformation and magmatism in the NCC. Widespread east –west Cretaceous extension represents the collapse of the Huabei collisional plateau (Zhang et al. 2001), and of the Yanshan belt in the northern NCC (Davis 2003), which led to the formation of the numerous metamorphic core complexes that are now widely recognized in the eastern North China Craton (Davis et al. 1996; Yang et al. 2004b; Cope & Graham 2007). These core complexes formed between 140 and 120 Ma (Cretaceous) and all seem to show a commonly oriented stretching lineation indicating extension or transport from NW to SE. Opening of the Bohai Sea (Allen et al. 1997) and many other marginal basins in the Tertiary shows that this extension was long-lived. Collision of India and Asia resulted in the uplift of numerous mountain ranges and

large-scale crustal thickening throughout Asia since about 50 Ma, and some of the young extension in Eastern Asia, including within the NCC, may be related to escape away from this collision (Molnar & Tapponnier 1975; Yin & Nie 1996).

Mesozoic to Cenozoic structural evolution and basin formation Many large Mesozoic and Cenozoic basins cover the eastern North China Craton (Fig. 11). The development of these large basins was concentrated in two time periods, Jurassic to Cretaceous and Cretaceous to present (Griffin et al. 1998). Ren et al. (2002) proposed that the overall NW– SE-trending extensional stress field was related to changes in convergence rates of India–Eurasia and Pacific – Eurasia combined with some asthenospheric upwelling. Sass & Lachenbruch (1979) assumed that the two stages of basin formation were related to lithosphere erosion that began in Early Jurassic times. However, some workers have related the extension to subduction of the Kula plate beneath Eastern China in Jurassic– Cretaceous times and later subduction of the Pacific plate (Griffin et al. 1998). Geophysical and geochemical data (Figs 12 and 13) show that the areas of thinner lithosphere correspond to the deepest Cenozoic basins (Yuan 1996; Griffin et al. 1998). Kimberlites found in these basins (Fig. 14) provide the only direct source of information about the underlying mantle. The Cretaceous –Tertiary Tieling basin in northern Liaoning Province (near Shenyang; Fig. 11) hosts Mesozoic– Tertiary kimberlites (Fig. 14; Griffin et al. 1998). Phanerozoic lithosphere beneath the Tan-Lu fault was replaced by hotter, more fertile material that may be related to the Tertiary rifting of the Shanxi highlands (Ke & Tian 1991; Dobbs et al. 1994; Zheng et al. 2001). Furthermore, the Eocene Luliang kimberlites imply that Phanerozoic-type mantle was in place by the end of the Cretaceous (Griffin et al. 1998). Another kimberlite within a narrow Cenozoic basin lying along the Tan-Lu fault in Tieling County (Fig. 14) shows similar Phanerozoic-type mantle that is related to rifting. Garnet temperatures at shallow depths indicate that significant cooling occurred after the Phanerozoic mantle was emplaced beneath this area (Griffin et al. 1998).

Cenozoic extension in the Shanxi graben and Bohai Sea basins Cenozoic extensional deformation in the central NCC is localized in two elongate graben systems surrounding the Ordos Block (Fig. 11): the S-shaped Weihe –Shanxi graben system (Shanxi

Fig. 10. Map of the Yanshan orogen, showing abundant normal faults and granitoid plutons.

OROGEN CRATON OROGEN CYCLE, NORTH CHINA 17

Kunlu n

fault

sh e uib

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uan faul Ximing t

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Fig. 11. Map of Northern China showing Cenozoic-active structures and basins in and around the North China Craton. Modified after Zhang, Y. Q. et al. (2003a).

100

an

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18 T. M. KUSKY ET AL.

OROGEN CRATON OROGEN CYCLE, NORTH CHINA

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Fig. 12. Map showing Bouguer gravity and the prominent north–south gravity lineament that strikes across China, crossing the NCC along the approximate boundary between thick lithosphere to the west and thin lithosphere to the east. The north–south gravity lineament is parallel to the Pacific subduction margin, perhaps suggesting a causal link.

grabens for short) to the east and SE, and the arc-shaped Yinchuan –Hetao graben system to the NW (Zhang, Y. Q. et al. 1998; Morley 2002). The southwestern margin of this block corresponds to a zone of compression (Zhang 1989), through which the North China Craton is in direct contact with the Tibetan Plateau (Yin & Harrison 2001). Wang & Zhang (1995) determined that the subsidence in these grabens began during the Eocene, and extended to the whole graben system during the Pliocene. The Shanxi graben system was the last to be initiated in Northern China, at about 6 Ma. These two extensional domains show differences in the thickness of the crust and lithosphere; the thickness changes sharply across the eastern edge of the Taihangshan Massif (Ma 1989) on the eastern side of the Shanxi graben system. Zhang, Y. Q. et al. (2003) showed that the Shanxi graben system consists of a series of en echelon depressions bounded by normal faults. Xu et al. (1993) noted the S-shaped geometry of the Shanxi

graben system, with two broad extensional domains in the north and south and a narrow transtensional zone in the middle. Both SPOT imagery interpretation and field analyses of active fault morphology show predominantly active normal faulting. Right-lateral strike-slip motion along faults that strike more northerly led Xu et al. (1993) to interpret the Shanxi graben system as a right-lateral transtensional shear zone, whereas Zhang et al. (1998, 2001) considered it to be an oblique divergent boundary between blocks within Northern China. Zhang, Y. Q. et al. (2003) suggested that NNE– SSW-oriented initial extension along the footwall of frontal range fault zones in northern Shanxi predates the Pliocene opening of the Shanxi graben and may be coincident with the Miocene Hannoba basalt flow (Figs 11 and 14). The direction of extension that prevailed during the initiation and evolution of the Shanxi graben system shows a northward clockwise rotation, from 300–3308 along its southern and middle portion to 330–3508

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T. M. KUSKY ET AL.

may have shared a common mechanism with that of the opening of the Japan Sea. First, the opening of the Japan Sea began at the end of the Oligocene around 28 Ma or earlier, and continued to the Middle Miocene, at about 18 Ma (Tamaki et al. 1992; Jolivet et al. 1990; Fournier et al. 1994); the youngest dredged basaltic volcanic rocks were dated at 11 Ma (Kaneoka et al. 1990). Second, the spreading direction of the Japan Sea is roughly north– south to NNE–SSW (Sato 1994), consistent with the Miocene stretching direction in Northern China. Finally, the same extensional stress regime trending ENE –WSW to NE –SW has been documented in northeastern Japan (east of the Japan Sea) based on the direction of dyke swarms and dated at 20 –15 Ma (Sato 1994).

Discussion: decratonization and the orogen to craton to orogen cycle

Fig. 13. Map showing depth to the low-velocity zone (modified after Griffin et al. 1998). NSGL, north– south gravity lineament.

across the northern part. SPOT imagery interpretation of late Quaternary active fault morphology by Zhang, Y. Q. et al. (1998) implies that the opening of the Shanxi graben system proceeded by northward propagation. This opening mode corroborates the kinematic interpretation by Zhang, Y. Q. et al. (2003a) and reflects a counterclockwise rotation of the Taihangshan Massif with respect to the Ordos Block around a pole located outside the block (Peltzer & Saucier 1996; Zhang, Y. Q. et al. 1998). During the Miocene, the regions of rifting in Northern China were subjected to regional subsidence and the eruption of widespread basalt flows (Fig. 14) Yang et al. 2006a, b. Basalt volcanism, dated by Liu et al. (1992) at 25– 10 Ma, was extensive in Mongolia and Eastern China, including the areas of the above grabens. According to Zhang, H. F. et al. (2003), this volcanism was related to extension in response to rollback of the subducted Pacific plate beneath Eastern Asia. Miocene normal faulting occurred particularly in the offshore part of the Bohai Sea basin, where this normal fault set strikes more easterly (Zhang, Y. Q. et al. 2003b). Liu et al. (2001) and Zhang, Y. Q. et al. (2003b) inferred that the Miocene extension in North China

Major north–south-striking topographic and gravity gradients that strike across the NCC (e.g. Liu 1992; Niu 2005) correspond to a major change in lithospheric structure (Fig. 12). The north–south gravity lineament is a major gradient in Bouguer gravity anomalies that corresponds roughly to the border between the Eastern and Western Blocks (or areas with and without root loss). It also, however, extends further north and south for thousands of kilometres beyond the borders of the NCC (Fig. 12). Because the gravity lineament also corresponds to areas of Tertiary basin formation along major faults, it may represent a major crustal structure parallel to the Pacific subduction zone. The north–south gravity lineament is also interesting because it bounds areas that to the west have thick crust and 150–200 km thick lithosphere (Fig. 13), large negative Bouguer anomalies, and low heat flow. Sub-Moho seismic Vp values west of the lineament are high, in the range of c. 8.1–8.3 km s21. However, to the east the crust and lithosphere are generally thinner, there is high heat flow, and the regional Bouguer anomalies are zero to slightly positive. Sub-Moho seismic velocities are lower than to the west, ranging from 7.6 to 7.7 km s21, with some faster regions (implying partial root loss?). Tomographic profiles from the Eastern Block (Yuan 1996) show an irregular velocity structure for the lower lithosphere, suggesting only partial root loss. The Eastern Block is seismically very active, experiencing many magnitude 8þ earthquakes that include six of the 10 most destructive events in recorded history (Kusky 2003), which killed more than one million people. From 3D P-wave velocity data Huang & Zhao (2004) established that in the lower crust and in the uppermost mantle under the source regions of the large earthquakes there

OROGEN CRATON OROGEN CYCLE, NORTH CHINA

21

Fig. 14. Map of the eastern NCC showing distribution of kimberlites of different ages that entrain up mantle xenoliths.

are low-velocity and high-conductivity anomalies, which they considered to be associated with fluids. The fluids caused weakening of the seismogenic layer, contributing to the initiation of the

large crustal earthquakes. These fluid data suggest that multiple subduction events beneath the zone of depleted lithosphere enriched the mantle in water, and hydro-weakened it. Whatever the

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process of root loss (e.g. Menzies et al. 1993; Griffin et al. 1998; Wilde et al. 2003; Wu et al. 2003a, b; Yang 2003; Deng et al. 2004; Fan & Menzies 1992a, b), it appears to have caused continuing lithospheric instability. Loss of the lithospheric root is also shown by the compositional data for mantle xenoliths brought up in early Palaeozoic and Mesozoic to Tertiary kimberlites and volcanic rocks (Fig. 14). The oldest kimberlites (490–450 Ma) are the Palaeozoic Fuxian and Mengyin pipes in the west, whereas the Tieling intrusions are Cretaceous to Tertiary in

112

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Mesozoic

age (Fig. 14). Xenoliths in basalts from Nushan are only 0.8 –0.5 Ma old, which, together with the older examples, provides a 500 Ma history of mantle samples from beneath the NCC. Geotherms based on mantle xenolith data (Ryan et al. 1996; Griffin et al. 1998; Xu et al. 1998) and garnet concentrates show that in Ordovician times, the Eastern Block had a low conductive cratonic geotherm, with many samples coming from beneath the diamond stability field. The Ordovician lithosphere –asthenosphere boundary is estimated to have been at about 180 km depth (Griffin et al.

Granites

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Yellow Sea Qin orog glong - Da en bi

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0

n

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Explanation Mesozoic granitoid

Mesozoic gold deposit

Fig. 15. Mesozoic gold and granite provinces of the NCC. (Note how the gold deposits and granites outline a ring around the Eastern block of the craton, suggesting that they may delineate the limits of the area of root loss). Modified after Goldfarb et al. (2001) and Wu et al. (2005).

OROGEN CRATON OROGEN CYCLE, NORTH CHINA

1998). In contrast, compositional data from the younger mantle samples reveal a high geotherm and a lithosphere–asthenosphere transition that had risen to about 80 km depth. Compositional data from xenoliths thus clearly show the loss of the lithospheric root beneath the eastern NCC, but do not yield information on exactly when this loss may have occurred, why it occurred, or what the loss means for cratonic evolution. Basalts erupted through the crust of the Eastern Block (Fig. 14) also show a change in composition from Mesozoic to Tertiary, with high-Mg andesites or adakites interpreted as evidence for lower crustal foundering in Jurassic – Cretaceous times. From geochemical and isotopic data for Mesozoic lavas of the eastern NCC, Zhang, H. F. et al. (2003)

23

concluded that there is thicker, less modified lithospheric mantle in the interior, and thinner, more heavily modified lithospheric mantle beneath the craton margins. They also demonstrated a secular change in the lithospheric mantle from a Palaeozoic refractory continental lithosphere to a Mesozoic enriched lithosphere. Although extending for thousands of kilometres along the Pacific rim, Mesozoic granitoids and gold deposits (Goldfarb et al. 2001; Hart et al. 2002; Mao et al. 2002; Wu et al. 2005) that are contemporaneous with the lithospheric thinning form a ring (Fig. 15) around the Eastern Block (Yang, J. H. et al. 2003). The removal of the lithospheric mantle and upwelling of new asthenospheric mantle induced partial melting and dehydration of the

Fig. 16. Model showing simplified evolution of the North China Craton, from orogen to craton to orogen, and how crustal and mantle root processes may be linked (note that the root is not to scale). Growth of the craton by subduction– accretion in arc settings probably involved the underplating of buoyant oceanic slabs (e.g. Kusky 1993), which would eventually become the subcontinental mantle root. Plume-influenced rifting at 2.7 Ga broke apart the future Eastern Block, and led to the development of a passive margin sequence on the western side of the Eastern block. This margin collided with a convergent margin at 2.5 Ga, amalgamating the craton. At 1.85 the craton experienced a major collision event along its northern margin, which resulted in partial replacement of the mantle root and widespread high-grade metamorphism, and the formation of a collisional plateau and foreland basin. For much of the Palaeozoic the craton was relatively internally stable, but accommodated about 18 000 km of cumulative subduction along its northern, southern, and eastern margins. Subduction-related dehydration reactions in the slab released fluids that hydrated the mantle, weakening its rheology and lowering its melting point, which allowed the root to release a low-density melt phase during Mesozoic tectonism, become denser, and sink into the asthenosphere after being triggered by near-simultaneous collisions along its northern (Solonker) and southern (Dabie–Sulu) margins. IMNHO, Inner Mongolia–Northern Hebei Orogen.

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lithospheric mantle and lower crust, and the derived fluids deposited the gold (Yang, J. H. et al. 2003, 2004). The granitoids and associated ore-bearing fluids may contain one of the best and most detailed records of the history of root loss beneath the NCC, perhaps preserving a history of the chemical and physical environments associated with foundering of subcrustal lithosphere. Additional research on these granitoids and mineral deposits may yield considerable insights into the physical and chemical processes associated with root loss. Many models and constraints have been proposed to explain the delamination of lithospheric roots in orogens, and we apply some of these models to loss of the lithospheric root beneath the NCC (see Fig. 16). Marotta et al. (1998) defined four major stages during a mantle ‘unrooting’ process: orogenic growth; initiation of gravitational instability until lithospheric failure; sinking of the detached lithosphere; relaxation of the system. Meissner & Mooney (1998) suggested that the basic driving force for delamination is the negative buoyancy of the continental lower crust and subcrustal lithosphere with respect to the warm, mobile asthenosphere. A likely cause of such negative buoyancy is a phase transformation in the lower crust from mafic granulite facies to eclogite (Morgan 1984; e.g. Kaban et al. 2003). Thus weakness in the lower crust during continental compression and extension is a key to the process of delamination. According to Schott & Schmeling (1998), full detachment of a delaminated lithospheric slab occurs only if the viscosity of the lower crust is greater than c. 1021 Pa s. Lithospheric roots or unsupported slabs of at least 100–170 km depth extent are needed to provide sufficient negative buoyancy to allow delamination and detachment. Gao et al. (1998a, b) applied geochemical data to the problem of delamination under the eastern NCC. They found that the lower crust in Eastern China contains c. 57% SiO2, which contrasts with the generally accepted models of mafic lower crust. They further suggested that eclogite from the Dabie –Sulu UHP belt is the most likely candidate as the delaminated material, and that a cumulative 37–82 km thick eclogitic lower crust is required to have been delaminated to explain the relative Eu, Sr and transition metal depletions in the crust of East – Central China. Delamination of eclogites can also explain the significantly higher than eclogite Poisson’s ratio in the present Dabie lower crust and upper mantle and the lack of eclogite in Cenozoic xenolith populations of the lower crust and upper mantle in Eastern China. However, considering that the lower crust contains c. 57% SiO2, and that xenoliths of lower crust in Cenozoic basalts in Hanuoba, North Hebei are garnet gabbro and two-pyroxene granulites,

Zhai et al. (2004) suggested that delamination of eclogites possibly occurred only at the northern and southern edges of the eastern North China Block. The thinning of the lithosphere could be related to thermal–chemical erosion with a mantle upwelling under the joint grip of the surrounding blocks, although its mechanism is not clear, and we favour the hydro-weakening mechanism discussed above (e.g. Niu 2005; Windley et al. 2005; Komiya & Maruyama 2006).

Conclusion The North China Craton has experienced one of the longest and most complex histories of any geological terrane on the planet (Fig. 16). Events from c. 3.5 to 2.7 Ga primarily reflect the extraction of melts from the mantle, probably in arc settings, and the amalgamation of many arcs to form some of the distinctive blocks in the craton. By 2.7 Ga the Eastern Block of the craton apparently was affected by a plume, associated with rifting of another block off the current western edge of the craton, which led to the opening of an ocean and deposition of a passive margin sequence on the western edge of the Eastern Block from 2.7 to 2.5 Ga. At 2.5 Ga the Eastern Block collided with a convergent margin now preserved in the Central Orogenic Belt, and apparently attached to the Western Block, obducting ophiolites and depositing a thick foreland basin sequence on the Eastern Block. This 2.5 Ga event culminated in the amalgamation of the North China Craton, and the intrusion of 2.4 Ga dykes and plutons across much of the central part of the craton. These Archaean –Palaeoproterozoic events are responsible for the initial formation of the root of the North China Craton, and we speculate that the first stages of root formation may have involved underplating of buoyant oceanic lithospheric slabs beneath convergent margins, as described by Kusky (1993). Interestingly, this mechanism would result in different parts of the subcontinental lithospheric mantle having different properties such as orientation of slabs (and internal olivine crystals), perhaps leading to a different susceptibility to delamination or root loss in the events later in the craton’s history. The craton experienced its strongest metamorphic event at 1.85 –1.8 Ga, related to continent–continent collision, which overprinted and obscured earlier events. Metamorphic evidence shows that the crustal thickness doubled, and pressures of metamorphism increase from south to north. Although the location of the collision has been disputed, sedimentological, structural, igneous, metamorphic and tectonic patterns clearly show that the collision was along the north margin of the craton (Fig. 16). This collision was so strong that

Closure of Dongwanzi ocean

Post-collision extension

Ocean opening, N margin craton

Major continent– continent collision, final amalgamation of NCC

Post-collision extension Quiescence?

Subduction under Dabie Shan

Subduction under Solonker

Indosinian Orogeny

Subduction under Pacific margin

2.4 Ga

2.4– 2.3 –2.1 Ga

1.9– 1.85 Ga

1.8 Ga 1800 –700 Ma

700 – 250 Ma

500 –250 Ma

270 –208 Ma

200 –100 Ma

Cratonic blocks form; remnants preserved Rifting then arcs active in Dongwanzi Ocean

Event

2.5 Ga

2.7; 2.55 –2.5 Ga

3.5–3.1 Ga

Age TTG, gneiss, fuchsitic quartzite, pelite Formation of TTG, CA arc suite, accretionary prisms, ophiolites, continental arc in Hengshan Formation, deformation, metamorphism of Central Orogenic Belt Formation of regional mafic dyke swarms, flood basalts Passive margin sediment, N margin; continental sediment, interior; collision of arcs at 2.3 and 2.1 Ga Formation of granulite plateau N part of craton, widespread metamorphism; deposition of S-prograding wedge of Changcheng clastic foreland basin Mafic dyke swarms Period of root stability, when craton acts like a craton; platform sediments? Cambro-Ordovician limestones deposited on platform, active margin on south 405 – 207 Ma, orogeny in NCC involves terrane accretion on N margin craton, and in Central Asia Orogenic Belt Scissor-like closure of Solonker Ocean Remelting of lower crust in Jiao-Liao massif

Signature in crust

Table 1. Summary of geological evolution of the North China Craton

Xiao et al. 2003

Hydration owing to ingress of slab fluids?

(Continued)

Li, S. Z. et al. 2006a

S. Z. Li, Hacker, Rathsburger, Rowley, Sengo¨r, Niu

Hydration owing to ingress of slab fluids?

Shortening of N edge of SCLM Hydration owing to ingress of slab fluids?

Peng, Li, Kusky, etc.

Gao et al. 2002, 2006; Kusky & Li 2003

Kusky & Li 2003; Kusky et al. 2006 Zhao, et al. 2002; Kusky & Li 2003; Zhao, T. P. et al. 2004

Kusky et al. 2001; Gao et al. 2002, 2006

Kusky et al. 2001; Li et al. 2002; Kusky 2004

Zhai et al. 2004

References

Mantle melting Stability

Possible collision-related delamination in part of craton? Replacement of part of root

Isotherm relaxation

Possible underplating of slabs beneath Western (+ Eastern) Blocks; mantle hydration Collision-related deformation of SCLM; formation of depleted root Melt extraction

Melt extraction

Signature in SCLM

OROGEN CRATON OROGEN CYCLE, NORTH CHINA 25

Regional extension

Adakites

Gold, etc. mineralization

Major volcanism

Himalayan Orogeny

140 – 105 Ma

160 –106 Ma

134 – 103 Ma

147 – 112 Ma

50 – O Ma

SCLM, subcontinental lithospheric mantle.

Present

165 – 90 Ma

Collision and post-collision thrusting in both Solonker and Dabie Shan collision zones Yanshanian Orogeny

Event

200 –150 Ma

Age

Table 1. Continued

Collision of India – Asia, uplift and exposure, extension Active normal faults, hot springs, volcanism in Eastern Block; quiet, low heat flow, no earthquakes in Western Block

Fluid flow on regional scales, gold mineralization

Formation of circum-Pacific magmatic belts, Tan-Lu fault Formation of many metamorphic core complexes, most have SE– NW extension directions; from 132 to 128 Ma, large-scale left-lateral motion on Tan-Lu fault (Zhu et al. 2001) A-type magmatic rocks

Thrust belts on craton margins, foreland sediments

Signature in crust

Extrusion

Overlaps with Yanshanian

Overlaps with Yanshanian

Overlaps with Yanshanian

Decompression?

Hydration owing to ingress of slab fluids

Thickening of crust– mantle system; loss of additional root?

Signature in SCLM

Liu et al. 2001; Zhang, Y. Q. et al. 2003a

Wei 2002; Xu et al. 2002; Davis 2003; Gao et al. 2006 Mao et al. 1999; Goldfarb et al. 2001; Yang et al. 2003 Zhang et al. 2000; Wang et al. 2001 Yin & Harrison 2001

Niu et al. 1994; Zhang 1989; Zhang et al. 1997; Zhang et al. 1998; Webb et al. 1999; Davis et al. 2002

Gao et al. 2002, 2004; Li, S.Z. et al. 2006a, b

References

26 T. M. KUSKY ET AL.

OROGEN CRATON OROGEN CYCLE, NORTH CHINA

in many places, particularly along the northeastern margin of the craton (Fig. 16), the 2.5 Ga subcontinental lithospheric mantle was apparently replaced by 1.8 Ga asthenosphere. For much of the Palaeozoic, the North China Craton was internally relatively stable, but c. 18 000 km of subduction along its northern, southern, and eventually its eastern margins led to extensive hydration of the mantle root, and preweakening of the root before massive Himalayanstyle collisions along the northern (Solonker) and southern (Dabie Shan) margins of the craton. These nearly simultaneous collisions in the Triassic strongly affected the mantle root, and when the upper crust entered a phase of orogenic collapse in the Jurassic and Cretaceous, the root appears to have similarly responded by somehow detaching and sinking into the asthenosphere, and/or being thermally eroded perhaps after the root was lost. Palaeopacific subduction also involved at least one episode of ridge subduction, and the role that the thermally pulse associated with this event may have played in the loss of the lithospheric root beneath the North China Craton has yet to be analysed. Analysis of the geological history of the craton thus clearly shows that crustal and mantle processes are linked, and that a better understanding of surface tectonic evolution can lead to a better understanding of the processes that trigger root formation, root loss, and decratonization. Recognition of the process of decratonization and the orogen to craton to orogen cycle in the North China Craton, which is still experiencing the terminal consequences of the loss of its root, lead us to consider how important this process may have been through Earth history. If the North China Craton has lost its root and essential properties of being a craton, is it possible that other cratons have been ‘decommissioned’ and incorporated into mountain belts as isolated fragments or terranes of Archaean rocks so common in younger orogens? If so, we may have to reconsider current models of continental growth. We thank our many colleagues who have worked with us in the North China Craton, and provided stimulating discussions about the interpretation of regional tectonics. We especially acknowledge the contributions of J. H. Li, S. Z. Li, A. Polat, L. Wang, A. Kroner, R. Rudnick, G. Zhou, G. Davis, F. Y. Wu, X. N. Huang, X. L. Niu, S. Cheng, G. Muzi, S. Wilde, W. J. Xiao and J. H. Guo. Reviews by R. Goldfarb, A. Polat, Y. L. Niu and S. Wilde greatly improved the manuscript. This work was supported by the Chinese Academy of Sciences grants KZCX1-07 awarded to M. G. Zhai and Rixiang Zhu, and US NSF grants 01-25925 and 02-07886 awarded to T.M.K., by St. Louis University, Peking University, University of Leicester, and Ocean University of China.

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Temporal and spatial distribution of Mesozoic mafic magmatism in the North China Craton and implications for secular lithospheric evolution H.-F. ZHANG State Key Laboratory of Lithospheric Evolution, Institute of Geology and Geophysics, Chinese Academy of Sciences, PO Box 9825, Beijing 100029, China (e-mail: [email protected]) Abstract: Mesozoic mafic magmatism in the North China Craton shows a clear temporal and spatial distribution. Mesozoic mafic volcanism occurred dominantly in the northern and southern margins of the craton, with episodic eruptions from Early Jurassic to Late Cretaceous time. In contrast, Mesozoic mafic magmatism, which produced gabbroic to dioritic intrusive complexes, occurred in the centre of the craton in areas such as the Taihang Mountains and the Luzhong region, and all the complexes were intruded at almost the same time in the Early Cretaceous. This temporal and spatial distribution of Mesozoic mafic magmatism shows a strong heterogeneity of the Late Mesozoic lithospheric mantle beneath the North China Craton and a secular evolution of the lithospheric mantle beneath it. The lithospheric mantle beneath the Luzhong region is slightly isotopically enriched; that beneath the Taihang Mountains has an EM1 character in Sr and Nd isotopic features (87Sr/86Sr)i ¼ 0.7050– 0.7066; 1Nd(t) ¼ 217 to 210); and it possesses EM2-like characteristics (87Sr/86Sr)i up to 0.7114) beneath the Luxi–Jiaodong region. The general enrichment suggests that the Mesozoic lithospheric mantle was distinctive compared with Palaeozoic and Cenozoic counterparts. The secular evolution of this variably enriched Mesozoic lithospheric mantle requires a considerable modification, transformation and reconstruction of the lithospheric mantle beneath the craton in Late Mesozoic time. The elemental and isotopic compositions and the coherence of the lithospheric changes with the formation of circum-craton orogenic mobile belts indicate that these rapid lithospheric changes and corresponding lithospheric thinning were tectonically related to the multiple subduction and subsequent collisions of circumcraton blocks. Dehydration melting of subducted oceanic and continental crustal materials produced silicic melts that migrated up and reacted with lithospheric peridotites to generate more fertile lithospheric mantle (‘wet’ low-Mg peridotites plus pyroxenite veins). This is demonstrated by the fact that beneath the southern and northern margins the mantle was strongly modified, but beneath the central craton the effects were less marked. Compositional mapping of olivine from mantle peridotitic xenoliths and xenocrysts entrained in Mesozoic and Cenozoic basalts and mafic rocks throughout the craton suggests a similar framework. The North China Craton provides convincing evidence that the nature of the refractory lithospheric mantle was considerably changed in chemical composition through time, and that the lithospheric destruction was triggered by multiple circum-craton subductions and collisions.

The North China Craton (NCC) has an evolutionary history distinctive from that of other Archaean cratons in the world; its eastern part (the Taihang Mountains and area to the east) became tectonically active in the Phanerozoic, as manifested by frequent earthquakes and magmatism. The occurrence of mid-Ordovician (465 Ma) diamondiferous kimberlites and their entrained garnet peridotitic xenoliths within the eastern NCC in the Mengyin and Fuxian kimberlite fields reveals the presence of old and cold Palaeozoic lithospheric mantle with a thickness greater than 200 km (Fan & Menzies 1992; Griffin et al. 1992, 1998; Menzies et al. 1993; Chi 1996; Menzies & Xu 1998). This was further demonstrated by investigations of mantle peridotitic xenoliths, xenocrysts and heavy mineral

concentrates, and supported by the geothermobarometry of solid mineral inclusions in diamonds and Re –Os isotopic data for peridotitic xenoliths entrained in these kimberlites (Harris et al. 1994; Meyer et al. 1994; Zheng 1999; Wang & Gasparik 2001; Xu 2001; Gao et al. 2002). The underlying lithospheric mantle was dominantly composed of refractory or major element depleted (alkalis, Ca, Fe, and Al) harzburgites and clinopyroxene-poor lherzolites (i.e. possessing the features of typical cratonic lithospheric mantle). However, systematic studies on the major and trace elements and Sr –Nd isotopic compositions of Cenozoic basalt-borne spinel lherzolite xenoliths show the existence of a thinner (,80 km) and hotter lithospheric mantle beneath the eastern NCC in the Cenozoic

From: ZHAI , M.-G., WINDLEY , B. F., KUSKY , T. M. & MENG , Q. R. (eds) Mesozoic Sub-Continental Lithospheric Thinning Under Eastern Asia. Geological Society, London, Special Publications, 280, 35–54. DOI: 10.1144/SP280.2 0305-8719/07/$15 # The Geological Society of London 2007.

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(Fan et al. 2000; Zheng et al. 2001), when the lithospheric mantle was fertile and isotopically depleted (i.e. having the features of ‘oceanic’ lithospheric mantle beneath the ocean basins or tectonically active orogens). This was further demonstrated by geophysical and palaeogeothermal data and Re – Os isotopic results (Ma 1987; Gao et al. 2002; Wu et al. 2003a; Xia et al. 2004). Therefore, it is suggested that large-scale lithospheric thinning beneath the eastern NCC took place between the Ordovician and the Cenozoic, leading to the loss of about 100 km of lithosphere (Fan & Menzies 1992; Menzies et al. 1993; Deng et al. 1994, 1996, 2004; Griffin et al. 1998; Menzies & Xu 1998; Fan et al. 2000; Xu 2001; Zheng et al. 2001). This also resulted in marked compositional changes of the lithospheric mantle from a cold, thick and refractory Palaeozoic lithosphere (Griffin et al. 1992; Menzies et al. 1993) to a hot, thin and fertile Cenozoic lithosphere (Fan et al. 2000). Now the questions are: When, where and how did such lithospheric thinning take place? How many times did the compositional changes of lithospheric mantle occur, once or twice? What was the geodynamic background responsible for the modification and destruction of the lithospheric mantle? Many recent investigations have been made of this important lithospheric thinning and many studies undertaken on the Mesozoic mantle-derived mafic rocks (Fan et al. 2001, 2004; Guo et al. 2001; Qiu et al. 2002; Zhang & Sun 2002; Zhang et al. 2002, 2003, 2004a, 2005; Chen & Zhai 2003; Guo et al. 2003; Yan et al. 2003; Zhai et al. 2003; Chen & Zhou 2004; Chen et al. 2004; Gao et al. 2004; Xu 2004; Xu et al. 2004a, b; Yang et al. 2004; Ying et al. 2004, 2006a; Lu et al. 2005; Wang et al. 2005; Zhou et al. 2005). This paper is a review based on the recognition of the temporal and spatial distribution of Mesozoic mafic magmatism throughout the craton. A systematic compilation of the geochemical data for these mafic rocks and compositional mapping of olivines from mantle peridotitic xenoliths and xenocrysts entrained from both Mesozoic and Cenozoic basalts are reported here, to probe the nature and composition of the lithospheric mantle beneath the craton and to understand the enrichment processes and the mechanism responsible for them.

Regional geology The NCC is bounded on the south by the Palaeozoic to Triassic Qinling–Dabie – Sulu orogenic belt (Li et al. 1993; Meng & Zhang 2000) and on the north by the Central Asian Orogenic Belt (Sengo¨r et al. 1993; Davis et al. 2001; Xiao et al. 2003; Fig. 1). The Qinling– Dabie– Sulu orogenic belt resulted from the continental collision between the NCC and

the Yangtze Craton in the Triassic (Li et al. 1993). The belt contains eclogites and a variety of gneisses that were derived from the Yangtze Craton (Xu et al. 1992; Liu et al. 2001). The finding of diamond and coesite in eclogites (Xu et al. 1992) and coesite inclusions in zircons from ortho- and para-gneisses (Liu et al. 2001) demonstrates that the upper and middle crustal units of the Yangtze Craton were subducted to ultrahigh-pressure depths (Liu et al. 2001). The presence of exsolution lamellae of clinopyroxene, rutile and apatite in garnets from eclogites suggests that the subduction of the Yangtze crust reached depths greater than 200 km (Ye et al. 2000). This subduction and subsequent collision may have significantly affected the lithospheric mantle beneath the southern NCC. To the north of the NCC, the generally east – west-trending Central Asian Orogenic Belt formed through north–south-directed subduction and arc –arc, arc –continent, and continent– continent collision mainly during the Palaeozoic (Sengo¨r et al. 1993; Davis et al. 2001; Buchan et al. 2002; Xiao et al. 2003), when multiple Ordovician to Permian oceanic arcs and the Mongolian microcontinent were amalgamated to the active margin of the NCC (Sengo¨r et al. 1993). The widespread occurrence of Palaeozoic ophiolites, Palaeozoic to Triassic blueschists, and synchronous igneous rocks suggests that the subduction commenced in the Early Palaeozoic and that the final collision occurred in the Permo-Triassic (Xu & Chen 1997; Robinson et al. 1999; Davis et al. 2001; Buchan et al. 2002; Xiao et al. 2003; Zhou et al. 2004). The Suolun (Solonker) suture and the Xilamulun Fault mark the southern boundary of this belt. The uplift of the Yanshan belt and the extensive Yanshanian magmatism in the northern NCC may be a consequence of this collision and post-collisional tectonics (Davis et al. 2001). The occurrence of high-Mg basaltic andesites on the northern margin of the NCC in the Late Mesozoic provides geochemical evidence for the existence of southward subduction just prior to the collision (Zhang et al. 2003). The NCC is composed of two Archaean continental blocks, the Eastern and Western Blocks, separated by a Proterozoic orogenic belt geographically corresponding to the Taihang Mountains (Zhao et al. 2002, and references therein). The craton has a basement of dominantly Archaean to Palaeoproterozoic TTG (tonalitic –trondhjemitic–granodioritic) gneisses and metavolcanic and metasedimentary rocks, covered by Sinian– Ordovician marine sedimentary rocks, Carboniferous –Permian continental clastic rocks, and Mesozoic basin deposits (Zhao et al. 2002). In the Western Block, the cratonic lithosphere has been stable since the Precambrian. However, extensive magmatism in the Eastern

DISTRIBUTION OF MESOZOIC MAFIC MAGMATISM

37

Fig. 1. Simplified tectonic map showing major tectonic units in eastern China and localities of Mesozoic basaltic fields and mafic intrusions on the NCC (modified from Zhang et al. 2004a). Two Archaean blocks in the NCC, the Eastern and Western Blocks, are separated by the Palaeoproterozoic orogenic belt. Basaltic fields (number in open square) or mafic intrusions (number in open circle): 1, Fangcheng basalt; 2, Mengyin shoshonite; 3, Laiwu gabbro– diorite complex; 4, Jinan gabbro; 5, Zouping gabbro–diorite complex; 6, Jimo basalt; 7, Jiaodong mafic dykes; 8, Donggang gabbro; 9, Guyi and Fushan gabbros; 10, Laiyuan gabbro– pyroxenite complex; 11–16, Western Beijing, Jianchang, Chifeng, Yixian, Fuxin, Zhangwu basalts; 17, North Huaiyang basaltic rocks; 18, Northern Dabie gabbro– pyroxenite complex; 19, Tongshi syenite; 20, Xinyang xenolith-bearing ultramafic volcanic pipe; 21, Early Tertiary basalt in Hefei basin; 22, Xishan basaltic rocks western to Beijing; 23, Datong basaltic rocks. Mesozoic lithospheric mantle blocks: A, Luzhong slightly enriched mantle; B, Taihangshan EM1 mantle; C, Luxi–Jiaodong EM2-like mantle; D, northern margin mixed mantle; E, Ordos old cratonic mantle. Dashed line shows the possible boundary between the isotopic regions.

Block is evidenced by Palaeozoic kimberlites, Mesozoic volcanic and plutonic rocks, and Cenozoic basalts (Fig. 1). Voluminous Jurassic – Cretaceous volcanic rocks were erupted into a series of small Mesozoic fault- and rift-related or subsidence basins in the craton (Fig. 1).

Temporal and spatial distribution of mafic rocks on the NCC Mesozoic mafic magmatism in the eastern NCC is distributed in four major tectonic blocks (Fig. 1): (1) the northern margin of the NCC (i.e. the Yinshan– Yanshan –Liaoxi regions); (2) the interior of the NCC or Luzhong region (mainly in the Luxi area, Shandong Province); (3) the southern margin of the NCC, including southwestern Luxi, Jiaodong and North Huaiyang; (4) the Taihangshan region.

Mesozoic volcanism was vigorous in the northern margin of the NCC, and mainly gave rise to the Nandaling, Diaojishan, Houcheng, Donglingtai and Donglanggou Formations west of Beijing, and the Nandaling, Diaojishan, Huaqiying and Zhangjiakou Formations in northern Hebei Province; these volcanic rocks erupted at 180–120 Ma (Luo & Li 1997; Zhou et al. 2001; Li et al. 2004; Lu et al. 2005). The Donglingtai and Zhangjiakou Formations are mainly composed of intermediate – acidic rocks and the others of mafic rocks. Moreover, Late Cretaceous (115 –120 Ma) mafic dyke swarms occur extensively in the Gubeikou– Nanyan region of western Beijing (Shao et al. 2001). There were four major periods of volcanism in the Liaoxi region (Chen et al. 1997; Wang et al. 2001; Zhang et al. 2003; Zhu et al. 2004a, b): Early Jurassic (Xinglonggou Formation), mid-Jurassic

38

H.-F. ZHANG

(Lanqi Formation), Early Cretaceous (Yixian Formation), and late Early Cretaceous (Fuxin Formation). Volcanic rocks of the Xinglonggou Formation are limited to the Tan-Lu Fault (Chen et al. 1997). The Lanqi Formation was more extensive, with several cycles of eruption, each beginning with high-Mg andesites and ending with trachytes and tuffs. The 1000– 2000 m thick Yixian Formation consists dominantly of andesites, trachyandesites, and trachytes (Zhang et al. 2003). The Fuxin Formation comprises scattered central-type volcanoes. The lavas of these volcanoes were erupted about 100 Ma ago, extensively eroded and primarily preserved as volcanic conduits such as those at Jianguo and Niutoushan (Zhang et al. 2003). In the Luzhong region in the interior of the NCC, several Mesozoic mafic –intermediate intrusive complexes contain gabbros and diorites (Zhang et al. 2004a). These include the Jinan gabbro complex (Guo et al. 2001), the Zuoping gabbro– diorite complex (Guo et al. 2003), and the Laiwu gabbro–diorite complex (Chen & Zhou 2004). Moreover, there are small amounts of mafic– intermediate volcanic rocks and carbonatites such as the Mengyin potassic porphyries (Qiu et al. 2002), the Jiyang basin basalts and porphyries, and the Zibo carbonatites (Ying et al. 2004). All these rocks were formed almost contemporaneously at 125 –115 Ma. In southwestern Shandong Province, the Jiaodong region and Dabie –Sulu orogenic belts in the southern margin of the NCC, voluminous mafic–intermediate magmatism mainly comprises intermediate –mafic intrusive complexes, and volcanic and alkaline rocks. In southwestern Shandong Province and the Jiaodong region, thick intermediate–mafic volcanic rocks of the Early Cretaceous Qingshan Formation (Fan et al. 2001; Guo et al. 2004; Ying et al. 2006a) include the 125 Ma Fangcheng basalts (Zhang et al. 2002), the Jimo bimodal volcanic rocks (Fan et al. 2001), and flood basalts of the Late Cretaceous Wangshi Group (Yan et al. 2003), accompanied by the Yinan gabbros (Xu et al. 2004a, b). Intermediate –alkaline intrusive complexes of 180– 190 Ma age (Xu W. L. et al. 2002; Xu Y. G. et al. 2004b) occur in the northern Jiangsu Province. Alkaline intrusive complexes such as the Early Jurassic Tongshi syenites and Early Cretaceous Longbaoshan syenites and monzonites occur in southwestern Shandong Province (Zhang et al. 2005). Ultramafic –mafic intrusive complexes occur in the northern Dabie region (Jahn et al. 1999), and basaltic trachyandesites of the Xiaotian Formation and mafic dykes are widespread in the northern Huaiyang region (Fan et al. 2004). Many contemporaneous porphyries of 121– 125 Ma have associated gold mineralization in these regions (Guo et al. 2004).

In the Taihang Mountains, Mesozoic magmatism lacks volcanism and comprises dominantly intermediate –mafic intrusive rocks. Laiyuan intermediate –mafic intrusive complexes of Late Jurassic age (Zhang et al. 2004a) occur in the northern Taihang Mountains. In the southern Taihang Mountains (i.e. Hebei – Shanxi–Henan boundary region), there are many Early Cretaceous intermediate –mafic to alkaline intrusive complexes, such as the Fushan, Xi’anli, Hongshan and Guzheng complexes (Xu & Lin 1990; Tan & Lin 1994; Luo et al. 1999; Chen et al. 2004; Zhang et al. 2005). Almost all the complexes were intruded around 125 Ma (zircon sensitive high-resolution in microprobe (SHRIMP) age, Peng et al. 2004), except for the Hongshan alkaline complex, which was intruded at 138 Ma (Zhang et al. 2005). Thus, Mesozoic mafic magmatism has a temporal and spatial distribution in the eastern NCC (Fig. 1). In space, intermediate –mafic volcanic rocks occur mainly in the southern and northern margins of the eastern NCC, whereas intermediate –mafic intrusive complexes crop out in the interior of the craton and the Taihang Mountains. In time, Mesozoic mafic– intermediate rocks were emplaced in five major periods: all Early Jurassic (180 –190 Ma) rocks occur in the southern and northern margins of the eastern NCC, whereas intermediate –mafic rocks are in the northern margin, and syenites and subalkaline diorite – monzonites in the southern margin. Magmatism in the period 130–160 Ma is rare; there are only intermediate –mafic volcanic rocks and high-Mg andesites in the northern margin, trachyandesites and syeno-monzonites in the southern margin, and gabbros in the northern Taihang Mountains. The main period of Mesozoic magmatism in the NCC was 110–130 Ma, and mafic –intermediate igneous rocks of this age occur extensively in all blocks of the eastern NCC. About 100 Ma alkali basalts are of asthenospheric origin in the northern margin of the eastern NCC (Zhang et al. 2003). Magmatic rocks of 65 –86 Ma occur mainly in the Jiaodong region, as represented by the Wangshi Group basalts and mafic dykes, some of which contain xenoliths of mantle peridotites, pyroxenites and granulites. The temporal and spatial characteristics of the Mesozoic intermediate – mafic magmatism in the eastern NCC have close relationships with the diverse temporal and spatial evolution of the underlying lithospheric mantle.

Petrology Mafic (SiO2 contents .52 wt%) volcanic rocks (Fig. 1) of the eastern NCC include the Fangcheng basalts (Zhang et al. 2002), the Mengyin shoshonites (Qiu et al. 2002), the Jimo basalts (Fan et al.

DISTRIBUTION OF MESOZOIC MAFIC MAGMATISM

2001), the Liaoxi basalts (Chen et al. 1997; Zhang et al. 2003; Lu et al. 2005; Zhou et al. 2005), the Luxi basaltic rocks of the Qingshan Formation (Ying et al. 2006a), the Xishan basalts (Li et al. 2004), and the Northern Dabie and North Huaiyang basalts (Fan et al. 2004; Wang et al. 2004). All these rocks are in the calc-alkaline series and are enriched in Si, Ca, Mg, K, Na, and large ion lithophile elements (LILE) and depleted in high field strength elements (HFSE), a geochemical signature similar to that of island arc volcanic rocks. Mafic intrusions such as the Jinan gabbros (Guo et al. 2001), Zouping and Laiwu gabbro–diorite complexes (Guo et al. 2003; Chen & Zhou 2005), Yinan gabbro–diorite complexes (Xu et al. 2004a, b), Hanxing gabbro–monzonite complexes (Tan & Lin 1994), Laiyuan gabbro–pyroxenite complexes (Zhang et al. 2004a) are also calcalkaline. All the mafic intrusions show similar chemical characteristics; some contain peridotite and/or pyroxenite xenoliths (Xu et al. 1999; Chen & Zhou 2005). Previous petrological and geochemical studies (Tan & Lin 1994; Chen & Zhou 2005) indicated that the gabbroic bodies did not undergo extensive fractional crystallization, but have compositions close to those of their original basaltic magmas. Therefore, the isotopic features of all these mafic rocks probably reflect the nature of their mantle source. Detailed descriptions of these individual rocks have been given in the studies cited above.

Sr – Nd – Pb isotopic compositions Mafic rocks from the eastern NCC have large variations in isotopic composition (87Sr/86Sr)i ¼ 0.7034–0.7114; 1Nd(t) ¼ 221 to 5; (206Pb/ 204 Pb)i ¼ 16.4 –18.2; D8/4 ¼ 230 to 130), and display a clear regional or spatial variation, especially in Sr–Nd isotopic compositions (Fig. 2). Gabbros from the Taihangshan region have relatively restricted Sr and Nd isotopic compositions (87Sr/86Sr)i ¼ 0.7050–0.7066; 1Nd(t) ¼ 217 to 210), typical of the EM1 signature. In contrast, basalts, gabbros and mafic dykes from the Luxi– Jiaodong region have extremely high (87Sr/86Sr)i ratios (up to 0.7114) with Nd isotopic compositions similar to those of the Taihangshan mafic intrusions, and are EM2-like (Fig. 2). This EM2-like isotopic signature of the rocks from the Luxi– Jiaodong region is also a feature of the basaltic rocks, gabbros and pyroxenites from the Northern Dabie area (Fig. 2). The latter were interpreted to have been derived from a mantle source that was strongly affected by crust –mantle interaction during deep subduction (Jahn et al. 1999; Fan et al. 2004).

39

In the Luzhong region the majority of gabbros have similar (87Sr/86Sr)i ratios, but higher 1Nd(t) values (87Sr/86Sr)i ¼ 0.7040–0.7065; 1Nd(t) ¼ 210 to 24), compared with those of gabbros from the Taihangshan region. A few gabbros from the Luzhong region have very low 143Nd/144Nd isotopic ratios (1Nd(t) , 215), but their Sr isotopic compositions are similar to those of the other gabbros. Basaltic rocks from the northern margin of the NCC exhibit large Sr and Nd isotopic variations (Fig. 2); they plot overwhelmingly within the mantle array (87Sr/86Sr)i ¼ 0.7034–0.7066; 1Nd(t) ¼ 29 to 5). All mafic rocks (basalts, basaltic rocks and gabbros) from the Luxi –Jiaodong, Luzhong and Taihangshan regions have similar Pb isotopic compositions (Fig. 3); that is low (206Pb/204Pb)i ratios (,17.6) and high D8/ 4(.50), different from basaltic rocks from the northern margin, which dominantly show higher (206Pb/204Pb)i ratios (.17.5) and low D8/4 values (230 to 60) (Fig. 3).

Discussion Mantle heterogeneity The lack of a positive correlation of (87Sr/86Sr)i with SiO2 or Mg-number in these Mesozoic mafic rocks (Fig. 4) suggests that the process of crustal contamination was insignificant in generating the major differences between the various sample groups. This is further supported by the high MgO contents of these rocks (Fan et al. 2001; Guo et al. 2001; Zhang et al. 2002, 2003, 2004a). The presence of mantle xenoliths in some basalts and intrusions (Xu et al. 1999; Zhang et al. 2002) demonstrates that the parental magmas ascended rapidly, so that significant contamination did not occur. Therefore, these mafic rocks may reflect the isotopic composition of the Mesozoic lithospheric mantle beneath the NCC. Data compiled in this study demonstrate a high mantle heterogeneity and a clear regional variation of the Mesozoic lithospheric mantle beneath the NCC. The lithospheric mantle beneath the central NCC (i.e. the Luzhong region) was slightly Sr–Nd isotopically enriched towards an EM1 component, that beneath the Taihang Mountains was characterized by an EM1 signature, and that under the Luxi–Jiaodong and Northern Dabie –North Huaiyang regions has an EM2-like character. In contrast, the lithospheric mantle beneath the northern margin appears to be more complicated, being generally depleted (DMM) but with some involvement of EM1 component (Fig. 2). The rocks from the northern margin show some evidence of magma mixing from different origins (lithospheric mantle and

40

H.-F. ZHANG

Fig. 2. Initial Sr –Nd isotopic diagram for Mesozoic mafic rocks (modified from Zhang et al. 2004a). Data sources are from the literature: Luzhong region (Guo et al. 2001; Liu et al. 2004; Zhang et al. 2004a; Ying et al. 2006a); Taihangshan region (Chen et al. 2004; Zhang et al. 2004a; Lu et al. 2005); Luxi– Jiaodong region including gabbros and mafic volcanic rocks from Northern Dabie and North Huaiyang areas (Fan et al. 2001, 2004; Qiu 2002; Zhang & Sun 2002; Zhang et al. 2002; Xu et al. 2004a, b; Yang et al. 2004; Wang et al. 2005; Ying et al. 2006a); northern margin of the NCC (Chen et al. 1997; Zhou et al. 2001; Zhang et al. 2003; Li et al. 2004; Lu et al. 2005). DMM, depleted mid-ocean ridge basalt (MORB) mantle; EM1, enriched mantle 1; EM2, enriched mantle 2. The data sources and symbols in the following figures are the same.

lower crust) as evidenced by the reverse zoning of clinopyroxene xenocrysts (Shao et al. 2006). Here, we name the above isotopically distinctive regions the Luzhong slightly enriched mantle, the Taihangshan EM1 mantle, the Luxi –Jiaodong EM2 mantle, and the northern margin ‘mixed’ mantle (Fig. 1). In contrast to the above regions in the Eastern Block, the Western Block lacks significant magmatism in the Mesozoic (Fig. 1). Therefore, the Mesozoic lithospheric mantle beneath the Western Block may have a different feature (i.e. inheritance from old and cold lithosphere).

Temporal – spatial evolution Available data clearly show that the Mesozoic lithospheric mantle beneath the NCC shows a spatial evolutionary trend towards a more radiogenic

Sr isotopic ratio and a less radiogenic Nd isotopic ratio from the centre (block A in Fig. 1) to the Luxi –Jiaodong (block C) and the Taihangshan region (block B). Interestingly, this mantle also shows a pronounced isotopic enrichment with time. In the Luxi– Jiaodong region, basaltic rocks ranging in age from 170 to 120 Ma become increasingly enriched in radiogenic Sr isotope compositions with time (Fig. 5), the highest (87Sr/86Sr)i ratio (0.712) being at 120 Ma. Similarly, in the northern margin of the NCC Mesozoic basalts show a similar evolutionary trend in Nd isotopes with time (Xu 2001). These significant temporal and spatial evolutionary trends were inevitably related to a significant tectonic and thermal event in eastern China. This temporal evolution was also elegantly demonstrated by the compositions of syenites and monzonites from different Mesozoic intrusive complexes (Zhang et al. 2005).

DISTRIBUTION OF MESOZOIC MAFIC MAGMATISM

Fig. 3. Initial 206Pb/204Pb v. D8/4 diagram for Mesozoic mafic rocks in the NCC (modified from Zhang et al. 2004a). D8/4 ¼ [(208Pb/204Pb)i 2 (208Pb/204Pb)NHRL]  100; (208Pb/204Pb)NHRL ¼ 1.209  (206Pb/204Pb)i þ 15.627. D8/4 . 0 indicates the presence of a high Th/Pb ratio, resulting in enrichment in radiogenic 208Pb.

Olivine compositional mapping Systematic compositional mapping of olivines from mantle peridotitic xenoliths and/or xenocrysts entrained in basaltic rocks and mafic dykes from the eastern NCC (Table 1 and Fig. 6) demonstrates that an apparently temporal and spatial variation

41

Fig. 4. SiO2 v. (87Sr/86Sr)i isotopic ratios for Mesozoic mafic rocks (modified from Zhang et al. 2004a).

and heterogeneity exists in the Mesozoic and Cenozoic lithospheric mantle beneath the eastern NCC (Fig. 7). This is entirely consistent with the temporal and spatial distribution (i.e. the block feature) of the Mesozoic lithospheric mantle beneath the NCC as revealed by the geochemistry of the lithosphere-derived mafic rocks discussed above. The different evolution of the Cenozoic

Fig. 5. Sr isotopic variation with time in Mesozoic and Cenozoic basaltic and shoshonitic rocks (modified from Zhang et al. 2004a).

42

H.-F. ZHANG

lithospheric mantle between the Taihang Mountains and the Tan-Lu Fault can be explained by the differences in nature and composition of the melt involved in the interaction with the old refractory lithospheric mantle. At the same time, olivine Fo mapping also reveals that the Tan-Lu Fault played an important role in both the Mesozoic and Cenozoic mafic magmatism and the lithospheric evolution in the eastern NCC. Furthermore, the olivine compositions indicate that some old lithospheric mantle remnants still exist in the eastern NCC after the lithospheric thinning. Thus, the wholesale delamination model for the lithosphere thinning (Wu et al. 2003b) in the eastern NCC is difficult to reconcile with the data. Detailed discussion on this section has been published (Zhang et al. 2006a).

Origin of mantle enrichment; interaction with silicic melt The above-mentioned mantle enrichment beneath the NCC can be interpreted as a result of influx of melts from a subducted continental slab, which over time developed isotopic heterogeneity (Zhang & Sun 2002; Zhang et al. 2002). If this is correct, the melt must be high in LILE and radiogenic Sr isotopic compositions, and low in HFSE and radiogenic Nd–Pb isotopic compositions. Silicic melts with such geochemical characteristics can only be derived from the partial melting of an old lower crust. High Th/U ratios of the Mesozoic magmatic rocks on the craton (enriched in 208Pb) indicate the possible involvement of middle, or even upper crust in their sources. The Mesozoic lithosphere undoubtedly evolved from Palaeozoic lithosphere, thus it should inherit some Palaeozoic lithosphere signatures. Mantle xenoliths have been discovered in Palaeozoic kimberlites from the NCC and these xenoliths have very restricted Nd isotopic compositions (1Nd  5; Zhang et al. 2002). Nd isotopic compositions of gabbros from the Luzhong region (1Nd(t) ,210) are slightly lower than those of the Palaeozoic kimberlite-borne mantle xenoliths from the NCC, suggesting that their mantle source had an inherited nature of old lithospheric mantle, with slight modification. In contrast, basalts from the Luxi–Jiaodong region have 1Nd(t) values (on average c. 216, Fig. 2) much lower than those of the Palaeozoic mantle xenoliths, indicating that the Mesozoic lithospheric mantle beneath the Luxi –Jiaodong region was severely modified by the addition of significant quantities of silicic melts. The reaction of silicic melts with the refractory Palaeozoic mantle peridotites may have produced pyroxenite veins in the mantle, as evidenced by the ubiquitous

occurrence of pyroxenite xenoliths in Fangcheng basalts (Zhang et al. 2002). This modified lithosphere (peridotite þ pyroxenite) became the source of Mesozoic calc-alkaline basalts and mafic intrusions. The crust-derived silicic melt required an early subduction process to bring the crust to mantle depth. Thus, it can be envisaged that the lithospheric mantle beneath the NCC experienced several evolutionary stages in the Phanerozoic, although subduction could be as early as Palaeozoic time. The Palaeozoic cratonic lithospheric mantle evolved into the Mesozoic enriched lithospheric mantle as a result of the interaction with a melt from the subduction of a continental slab, but this Mesozoic mantle did not survive for a long time period, because the Cenozoic lithospheric mantle underneath the NCC again shows contrasting features; that is, it is now fertile but depleted in Sr– Nd isotopic compositions (Fan et al. 2000).

Timing of the lithospheric thinning The exact timing of the lithospheric thinning of the eastern NCC (i.e. the start, accomplishment and peak of lithospheric thinning) remains hotly debated (Zhai et al. 2003, and references therein). The occurrence of mid-Ordovician (c. 465 Ma) diamondiferous kimberlites in the Mengyin and Fuxian kimberlite fields (Chi 1996) indicates the onset of instability of the cratonic lithosphere beneath the eastern NCC. These kimberlites were the first phase of magmatism since the Precambrian. Radiogenic isotopic and platinum group element (PGE) geochemistry of these kimberlites indicates that this ultramafic magmatism was related to the northward subduction of the Palaeo-Tethyan plate and/ or the southward subduction of the Palaeo-Asian plates (Zhang et al. 2003; Zhang et al. unpubl. data). The violent collision between the NCC and the Yangtze Craton in the Triassic not only produced one of the largest continent– continent collisional orogens in the world (i.e. Dabie –Sulu high –ultrahigh-pressure metamorphic belt), but also had a significant effect on the integrity and structural regime of the lithospheric mantle beneath the southern NCC. This continent–continent collision destroyed the stability of the cratonic lithosphere beneath the southern NCC, perhaps implying the actual commencement of the Mesozoic lithospheric thinning in the region. Early Jurassic (190 –180 Ma) alkaline magmatism in the southern margin of the craton, represented by the intrusive Tongshi syenite complex, was the first magmatic activity after the collision (Zhang et al. 2005). The occurrence of these alkaline rocks with the geochemical characteristics of an asthenospheric origin (Zhang et al. 2005) illustrates that at least part of the lithosphere

Locality

Host rock

Age of host rock

Mafic volcanic breccia

Zhucheng

Ju¨nan

4

5

Peridotitic xenolith Peridotitic xenolith

86.0 + 1.6 Ma

66.8 + 1.5 Ma

Late Cretaceous

Peridotitic xenolith Peridotitic xenolith

Olivine xenocryst

119.1 + 2.3 Ma

73.5 + 0.3 Ma

Olivine xenocryst

124.9 + 1.8 Ma

Type of olivine

Chifeng Pingzhuang

Alkali basalt

Alkali basalt

Fuxin

Fuxin Jianguo

Fuxin Wuhuanchi

Volcanic rocks Yixian Formation Alkali basalt Fuxin Formation Alkali basalt

Fuxin

107.3 + 1.6 Ma

8

Minggang

Mafic volcanic breccia

178.3 + 3.8 Ma

Peridotitic xenolith

Peridotitic xenolith Peridotitic xenolith Peridotitic xenolith

92.1 + 2.1 Ma 84.8 + 1.6 Ma

Peridotitic xenolith

Olivine xenocryst

105.5 Ma

122.7 Ma

Mesozoic: Zhumadian region (southern margin of the NCC)

7

6

Mesozoic: Liaoxi Fuxin and Inner Mongolia Chifeng regions (northern margin of the NCC)

Flood basalt Wangshi Group Mafic dykes

Daxizhuang

3

Volcanic rocks Qingshan Formation Mafic dykes

Huimingzhuang

Pishikou

Tholeiite Qingshan Formation

Fangcheng

2

1

Mesozoic: Luxi and Jiaodong regions (southern margin of the NCC)

Position

91.5– 89

91.5–86.4

93– 89

91.5– 89

90.5– 75.6

92– 88

87.4– 86.9

89– 86

92.2– 88

93– 90

91.9– 76

Fo range

Olivine was altered to talc

91.5

91.5

93

91.5

90.5

92

89

92.2

93

91.9

Fo maximum

89

86.4

89

89

75.6

88

86

88

90

76

Fo minimum

Lu et al. (2003) (Continued)

Zhang et al. (2003); this study

Zheng et al. (1999)

Wang et al. (2002)

Xu et al. (1999); Zhang et al. (2003); Zhu et al. (2004a)

Zhu et al. (2004b); Shao et al. (2005)

Ying et al. (2006b)

Han & Fu (1993)

Yan et al. (2003); Zhang et al. (2006b)

Zhang J. et al. (unpubl. data)

Pei et al. (2004)

Zhang et al. (2002, 2004b); Zhang (2005)

Data source

Table 1. Olivines from mantle peridotitic xenoliths or xenocrysts entrained in Mesozoic and Cenozoic basalts and basaltic rocks from the eastern North China Craton

DISTRIBUTION OF MESOZOIC MAFIC MAGMATISM 43

Locality

Host rock

Age of host rock

Hannuoba

Fansi

Xiyang –Pingding

Xuehuashan

Hebi

10

11

12

13

14

Alkali Ol basalt Alkali Ol basalt Alkali basalt

Alkali basalt

Alkali basalt

Alkali basalt

Late Tertiary

Late Tertiary

7.8 Ma

Late Tertiary

Tertiary

Tertiary

Alkali basalt Alkali basalt Alkali basalt Alkali basalt Basanite

Changle

Shanwang

Beiyan

Lingqu

Nu¨shan

Kuandian

18

19

20

Late Tertiary

0.55– 0.72 Ma

Late Tertiary

Late Tertiary

Late Tertiary

Late Tertiary

Late Tertiary

5.6– 5.7 Ma

Fo ¼ 100 Mg/(Mg þ Fe) (mol%). (For location of position numbers, see Fig. 6.)

Alkali basalt

Alkali basalt

Qixa

17

Alkali basalt

Penglai

16

Cenozoic: along the Tancheng – Lujiang wrench fault zone 15 Wudi Alkali basalt 0.73 Ma

Pingquan

9

Cenozoic: Taihang Mountains and Northern Hebei region

Position

Table 1. Continued

Peridotitic xenolith Peridotitic xenolith Peridotitic xenolith Peridotitic xenolith Peridotitic xenolith Peridotitic xenolith Peridotitic xenolith Garnet lherzolite Spinel lherzolite Peridotitic xenolith

Peridotitic xenolith Olivine xenocryst Olivine xenocryst Peridotitic xenolith

Peridotitic xenolith Peridotitic xenolith

Type of olivine

93.7– 88.3

92.2– 88.3

92.2 93.7

90.2– 82.2

90– 83

90– 83

91– 83

91– 83

92– 89.5

90.8– 89.5

90

93– 88

92.3– 84.5

92.3– 85.2

91.7– 86

92– 89.5

92.5– 89.4

Fo range

90.2

90

90

91

91

92

90.8

93

92.3

92.3

91.7

92

92.5

Fo maximum

88.3

88.3

82.2

86

83

83

86

89.5

89.5

88

84.5

85.2

86

89.5

89.4

Fo minimum

E & Zhao (1987); Liu (1992)

Xu X. S. et al. (1998)

This study

Wang et al. (1987); Zheng et al. (1998) Wang & Li (1996); this study

This study

Chen & Peng (1985); Gao & Zhang (1994) Chen & Peng (1985); Fan et al. (2000) Rudnick et al. (2004)

Zheng et al. (2001); this study

Tang et al. (2004)

Fan & Hooper (1989); Song & Frey (1989); Rudnick et al. (2004) Fan et al. (2000); Tang et al. (unpubl. data) Tang et al. (2004)

E & Zhao (1987); Liu (1992)

Data source

44 H.-F. ZHANG

DISTRIBUTION OF MESOZOIC MAFIC MAGMATISM

45

Fig. 6. Distribution of Mesozoic (number in open square) and Cenozoic (number in open circle) mantle peridotitic xenoliths-or olivine xenocryst-bearing basalts or basaltic rocks. 1, Luxi Feixian; 2, Jiaodong Qingdao; 3, Jiaodong Jiaozhou; 4, Jiaodong Zhucheng; 5, Jiaodong Ju¨nan; 6, Liaoxi Fuxin; 7, Inner Mongolia Chifeng; 8, Henan Zhumadian; 9, Hebei Pingquan; 10, Hebei Hannuoba; 11, Shanxi Fansi; 12, Shanxi Xiyang– Pingding; 13, Hebei Jingxing; 14, Henan Hebi; 15, Shandong Wudi; 16, Shandong Penglai; 17, Shandong Qixia; 18, Shandong Changle– Lingqu; 19, Anhui Nu¨shan; 20, Liaodong Kuandian.

had thinned to a thickness of 80 –100 km, implying that the commencement of the lithosphere thinning was locally earlier than the Early Jurassic. The peak of magmatism (intrusive and extrusive) was in the Early Cretaceous (130 –110 Ma), which was also the peak of the lithospheric thinning in the eastern NCC (Zhai et al. 2003). Magmatic activity later subsided after about 25 Ma (Xu 2001; Zhang et al. 2004a). The occurrence of the Late Cretaceous (86– 65 Ma) Wangshi Group flood basalts and mafic dykes, which have the geochemical characteristics of an asthenospheric source (Yan et al. 2003; Ying et al. 2006b), demonstrates that the lithospheric thinning had basically ceased by then, and the thickness of the lithosphere reached its thinnest, about 65 km (Xu 2001). The xenoliths record that the lithosphere beneath the region was accreted since Late Cretaceous time (Ying et al. 2006b) and continued to become thicker during Cenozoic time (Fan et al. 2000); the present thickness of the lithosphere is about 80 km. It should be noted that the process of lithospheric thinning beneath the eastern NCC seems to have been earlier in the north than in the south of the craton;

for example, large-scale calc-alkaline volcanism in the northern margin occurred in the Early Jurassic (the Nandaling Formation), apparently earlier than large-scale calc-alkaline volcanism in the southern margin (the Qingshan Formation, Zhang et al. 2003; Li et al. 2004). Similarly, asthenosphere-derived basaltic eruptions in the northern margin, represented by the Jianguo alkali basalts, were also much earlier than those in the southern margin (Zhang et al. 2003).

Mechanism for the lithosphere thinning The mechanism of lithosphere thinning has recently been hotly debated. The various models suggested for the lithosphere thinning can be generally classified into two groups: (1) a ‘delamination’ model (Wu et al. 2003b; Gao et al. 2004); (2) a ‘thermal erosion’ model, including thermal –chemical and thermal –mechanical (Xu 2001; Lin et al. 2004), mantle replacement (Zheng 1999; Zheng et al. 2001) and ‘mushroom’ models (Lu et al. 2000). The delamination of mafic lower crust and underlying lithospheric mantle could occur in orogenic

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Fig. 7. Olivine Fo % variation. Olivines from kimberlites in the NCC are from Zheng et al. (2001). Data for olivines from Mesozoic pyroxenite xenoliths and olivine phenocrysts in basalts are from Zhang (2005).

belts where thickened crust gave rise, for example, to the high-Mg adakitic volcanic rocks of the Xinglonggou Formation on the northern margin of the NCC, which could be the product of delaminated lithosphere (Gao et al. 2004). However, there is no evidence for lithosphere delamination in the central and southern NCC, because lithosphere delamination is a relatively rapid event, and abrupt delamination inevitably results in rapid upwelling of the asthenosphere. Experimental petrology shows that decompressional melting of upwelling asthenosphere in this environment would produce large-volume basalts with geochemical characteristics similar to those of the asthenosphere (i.e. similar to those of continental rift settings). However, major eruption of asthenosphere-derived basalts did not take place in the craton in the

Early Cretaceous, which is widely considered to be the time of lithosphere delamination (Wu et al. 2003b), but in Late Cretaceous and Cenozoic times. Furthermore, the relatively long and continuous Mesozoic magmatism in the eastern NCC (190 –110 Ma) is also difficult to reconcile with the delamination model. Extensive Early Cretaceous granitoid intrusions in the eastern NCC were considered to be the result of direct contact of the asthenosphere with the lower crust after the delamination (Wu et al. 2003b). However, in this period large-scale granitoid magmatism was widespread in NE and SE China (Zhou & Li 2000; Wu et al. 2005), away from the NCC. Voluminous Mesozoic granitoid magmatism in eastern China was probably due to a large heat supply, so it is critical to study the cause of a regional thermal anomaly in eastern

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Asia in the Early Cretaceous. This regional thermal anomaly could be related to the Kerguelen plume activated in the Pacific and/or Indian Ocean at 130 –120 Ma (Coffin et al. 2002), or to the break-up of the Pangaea supercontinent. Therefore, the mechanism for Mesozoic lithosphere thinning in the interior of the NCC probably was mainly thermal erosion with interaction between old lithospheric mantle and infiltrated melt. This is consistent with the spatial and temporal heterogeneity of lithosphere thinning in the eastern NCC, the relatively long interval of magmatism and dominantly calc-alkaline Mesozoic volcanism, and the lack of large-scale asthenospheric basalts. Extensional lithosphere thinning. It should be noted that mechanically extensional thinning of the lithosphere resulting from heating from below could be another important means of lithosphere thinning. Lithosphere extension and lower crust flux initiated by stress relaxation after the collision of the NCC with the Yangtze Craton and the Mongolian and Siberian plates led not only to the formation of various small Jurassic (and Early Cretaceous) sedimentary basins (Lin et al. 2005), but also to a decrease in lithosphere thickness. Back-arc expansion resulting from the northwestern subduction of the Pacific plate since the Cretaceous played a pivotal role in the extensional lithosphere thinning and the generation of Late Cretaceous and Cenozoic sedimentary basins in eastern China (Ren et al. 2002). An excellent example of such extensional lithosphere thinning is in the Luxi and Jiaodong regions. As mentioned above, the estimated thickness of the lithosphere in the region in the Early Cretaceous (130 –110 Ma) was about 80– 100 km (Zhang et al. 2005) and in the Late Cretaceous (86–65 Ma) was only about 65 km (Xu 2001; Ying et al. 2006b). There was no magmatism between these periods. Petrological and geochemical studies as well as geothermobarometry of spinel lherzolites of remnant lithospheric mantle entrained in Late Cretaceous basaltic breccias at Ju¨nan, Shandong Province (Ying et al. 2006b) indicate that some high Mg-number peridotitic xenoliths were once equilibrated in the garnet stability facies and when entrapped as xenoliths were relocated in the spinel stability facies. This provides direct evidence for extensional lithosphere thinning.

Geodynamics of lithosphere thinning and tectonic model There is a recent general consensus that the Mesozoic lithosphere thinning and modification was not an isolated event, but was closely related to the

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formation and evolution of the orogenic belts surrounding the NCC (i.e. result of its interaction with the circum-craton tectonic blocks). Now the question is which block was the most involved and how large was the effect? Subduction and collision of the Yangtze Craton. Northward subduction of the PalaeoTethyan Ocean since the Late Palaeozoic and subsequent subduction of the Yangtze Craton beneath the NCC as well as the collision between the NCC and the Yangtze Craton in the Triassic triggered instability of the lithosphere beneath the southern NCC. Thermal perturbations produced by local mantle convection above the subduction zone could have played an important role in the thermal erosion of the diamond- and garnet-facies mantle at the base of the lithosphere, leading to the removal of the diamond-facies and most of the garnet-facies lithospheric mantle. Silicic melts derived from melting of the subducted Yangtze crustal materials migrated up into the overlying lithospheric mantle, which could have resulted in a dramatic compositional change of the lithospheric mantle through peridotite –melt interaction. This process has been recorded in the geochemical features of the Mesozoic volcanic rocks (Zhang et al. 2002). Alkaline intrusive complexes, occurring in a zone broadly parallel to the Qinling –Dabie orogenic belts in the southern margin of the NCC, were considered by Yan et al. (2001) to be the direct product of post-collisional magmatism; they provide the geological evidence for the effects of Yangtze subduction and collision on the lithosphere beneath the southern NCC. Large differences in the isotopic composition of Mesozoic lithospherederived magmatism between the various blocks (Fig. 2) suggest that this subduction and subsequent collision had a significant effect on the evolution of the lithosphere beneath the southern NCC. Geochemical studies indicate that the effect could extend to the Laiwu–Tai’an line, about 450 km from the Qinling –Dabie orogenic belt (Zhang & Sun 2002; Zhang et al. 2003). Thus, the northward subduction of the Yangtze Craton and collision with the NCC was the force that triggered the lithosphere thinning and compositional change of the lithospheric mantle beneath the southern region of the NCC. Closure of the Palaeo-Asian ocean. The northward or southward subduction of the Palaeo-Asian ocean since the Early Palaeozoic (Sengo¨r et al. 1993) and the closure of the Late Palaeozoic Central Asian ocean and the Mongol–Okhotsk Sea not only resulted in the formation of the East Asian accretionary orogenic belts, but may also have had an important effect on the lithospheric

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Fig. 8. Schematic sections through the NCC illustrating reactivation of the eastern NCC, transformation of old cratonic lithosphere to Mesozoic fertile lithosphere, and lithospheric thinning as a result of post-orogenic lithospheric extension and eruption of magma (modified from Zhang et al. 2003). (a) Palaeozoic northward subduction of the Proto-Tethyan ocean and southward subduction of the Mongolian ocean beneath well-stratified cratonic lithosphere induced dehydration melting of the subducted crust (hydrothermally altered basalts and terrigenous sediments), which produced high-K melts. Migration of these melts initiated partial melting of the lowermost lithosphere and the uppermost asthenosphere to produce kimberlites. (b) Triassic collision of the Yangtze Craton and the Mongolian microcontinent with the NCC triggered destruction of the NCC lithosphere, simultaneously with formation of the Dabie Orogenic Belt, the Central Asia Orogenic Belt, and the Yanshan fold and thrust belt. Tectonic underplating of the Yangtze and Mongolian crust beneath the lithosphere of the NCC uplifted both sides of the NCC. (c) After 30–50 Ma the subducted continental slab produced silicic melts, which modified the overlying refractory lithospheric mantle. Continuous

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structure and evolution of the northern margin of the NCC. Dehydration of subducted oceanic lithosphere could have metasomatized the overlying mantle wedge, forming ‘wet’ peridotites, in turn leading to a change in the nature and composition of the lithosphere. Partial melting of a wet refractory lithospheric keel generated typical high-Mg basaltic andesites (Zhang et al. 2003), adakites and volcanic rocks that incorporated material from the oceanic crust (Li et al. 2004) on the northern margin of the NCC. Alkaline complexes were intruded in a zone nearly parallel to the Central Asian Orogenic Belt in the Late Triassic and Early Jurassic, and were considered by Yan et al. (2001) to be the direct product of the postcollisional magmatism. Therefore, the closure process could be an important force that significantly affected the lithospheric evolution beneath the northern margin of NCC. Differences in lithosphere thinning and modification beneath the southern and northern margins of the craton could be the result of differences in the formation and evolution of the southern and northern orogenic belts surrounding the craton. At the northern margin subduction of the Palaeo-Asian ocean was mainly of oceanic crust, and there was soft collision between accretionary wedges. The collisional force was relatively small and the contribution of the subducted ocean to the lithospheric mantle beneath the northern margin was dominantly fluid (i.e. there was no large-scale melt addition). Thus, the degree of the lithosphere enrichment beneath the region was relatively low, as demonstrated by the source of the high-Mg basaltic andesites (Zhang et al. 2003). In contrast, at the southern margin the Dabie –Sulu orogenic belts were mainly formed by deep subduction of continental lithosphere and by continent– continent collision. The contribution of subducted continental lithosphere to the lithospheric mantle beneath the southern margin was silicic melt addition. Therefore, the lithosphere enrichment beneath the southern margin was much greater (Zhang et al. 2004a). Subduction of the Pacific Ocean and heat source. Although Chen et al. (2004, 2005) argued

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that Mesozoic magmatism in the eastern NCC, even including the Taihang Mountains, resulted from westward subduction of the Palaeo-Pacific plate, they also suggested that this magmatism was the result of magma mixing of melt from the subducted plate and lithospheric melt. However, the data seem to show that there was no apparent material contribution from the Palaeo-Pacific plate to the Mesozoic magmatism in the eastern NCC (Fan et al. 2001; Guo et al. 2001; Xu 2001; Qiu et al. 2002; Zhang et al. 2002, 2004a, 2005; Xu et al. 2004a, b; Ying et al. 2006a). Thus, the effects of PalaeoPacific plate subduction on the eastern NCC were probably in the form of lithospheric extension and the production of a thermal anomaly. Back-arc expansion arising from the subduction (Ren et al. 2002) resulted in further extension of the Late Mesozoic lithosphere beneath the eastern NCC. Subduction of the Pacific plate may have provided the thermal source for the Mesozoic magmatism in the eastern NCC (Ren et al. 2002). The presence of this thermal anomaly triggered the partial melting of the already thinned and considerably modified ‘wet’ lithospheric mantle, which produced the extensive and disseminated calc-alkaline magmatism. Therefore, Palaeo-Pacific plate subduction was not the driving force for the large-scale lithospheric thinning and compositional modification in the eastern NCC, but it probably provided the thermal source for the Mesozoic magmatism and the extensional thinning of the Mesozoic lithosphere. Tectonic underplating model. Based on these observations, a tectonic underplating model of circum-craton subduction and collision is proposed for the lithospheric thinning and compositional transformation, as shown in Figure 8. Details of the events are given in the figure caption.

Conclusions Mesozoic basalts and mafic intrusions in the NCC were generally uncontaminated by continental crust, and therefore can be used to probe the nature of their mantle sources. The Mesozoic

Fig. 8. (Continued) modification by these melts converted the old cratonic lithospheric mantle to Mesozoic-enriched lithospheric mantle. Post-orogenic lithospheric extension caused decompressional melting of the enriched lithosphere during Jurassic–Cretaceous time, producing high-Mg basaltic andesites and high-K calc-alkaline volcanic rocks. Basaltic magma was underplated at the base of the crust, forming newly accreted lower crust. (d) Quantitative Late Cretaceous magmatism and back-arc expansion triggered by the Pacific plate subduction resulted in further lithosphere thinning beneath the eastern NCC. Subsequently, the uppermost asthenosphere converted to lithosphere with a decrease in geotherm gradient (i.e. forming newly accreted fertile lithospheric mantle of dominantly spinel lherzolite). Accretion of the lithosphere began in the Late Cretaceous and continued to the present. The lithosphere since the Late Mesozoic comprised the remnant Mesozoic lithospheric mantle in the upper part and the newly accreted lithospheric mantle in the lower part. The presence of slightly enriched mafic rocks in the centre and highly isotopically enriched rocks on both sides of the NCC supports this geodynamic model.

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lithospheric mantle beneath the NCC was highly heterogeneous in both isotopic compositions and olivine Fo values. The mantle beneath the central craton and the Taihangshan region had EM1 features, whereas the lithospheric mantle beneath the Luxi–Jiaodong region was EM2-like. The Mesozoic lithospheric mantle underneath the NCC can be envisaged as having evolved from Palaeozoic mantle that was severely modified by a silicic melt from subducted continental and oceanic slabs. The enrichment and transformation from a refractory to a fertile lithospheric mantle was produced by mantle peridotite–melt interaction. Thus the mechanism and driving force that triggered the lithospheric thinning and compositional modification of the NCC in the Mesozoic possibly resulted from the circumcraton subduction and subsequent collisions. This work was funded by grants from the Chinese Academy of Sciences (KZCX1-07) and the Natural Science Foundation of China (40225009, 40534022). B. Windley, M. Le Bas, and M. Sun are thanked for their careful reviews.

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Contributions of the lower crust to Mesozoic mantle-derived mafic rocks from the North China Craton: implications for lithospheric thinning F. HUANG1,2, S.-G. LI1,* & W. YANG1 1

CAS Key Laboratory of Crust – Mantle Materials and Environments, School of Earth and Space Sciences, University of Science and Technology of China, Hefei, Anhui 230026, China (*Corresponding author e-mail: [email protected]) 2

Department of Geology, University of Illinois at Urbana– Champaign, 1301 W. Green Street, IL 61801, USA

Abstract: The lithospheric mantle underneath the North China Craton changed completely from the Palaeozoic to the Cenozoic. This study reviews geochemical data from Mesozoic mantlederived mafic rocks from the North China Craton to investigate the role of mafic lower continental crust in lithosphere replacement. Samples from the North China Craton have typical ‘continental’ geochemical signatures, including depletion of high field strength elements, enrichment of large ion lithophile elements and Pb, unradiogenic Pb isotopes, and enriched Sr–Nd isotopic ratios. Positive correlation between initial 87Sr/86Sr and 206Pb/204Pb, low Ce/Pb and Nb/U, high Ba/ Nb and La/Nb, and unradiogenic Pb isotopes of Mesozoic mafic rocks cannot simply be explained by derivation from a lithospheric mantle enriched by ancient (Archaean or Mesoproterozoic) fluid or melt metasomatism. Instead, they more probably result from a lithospheric mantle or upwelling asthenosphere underneath the North China Craton that was modified by the lower continental crust in the Mesozoic. Because oceanic plate subduction zones surrounded the North China Craton during the late Palaeozoic, the lithospheric mantle underneath the North China Craton was weakened by fluids derived from subducted slabs, and thus shortened and thickened by continent –continent collisions of the North China Block with the South China Block and the Siberian plate. Metamorphic reactions occurred in the mafic lower continental crust beneath the North China Craton, creating garnet-bearing assemblages (eclogite and garnet pyroxenite) with densities of up to 3.8 g cm23, which led to negative buoyancy in the over-thickened lithosphere. The unstable lithosphere was delaminated and subsided into the uppermost mantle. The delaminated lower crust partially melted, producing SiO2-rich melts that metasomatized surrounding asthenospheric mantle, which upwelled and replaced the volume formerly occupied by the delaminated lithospheric mantle, resulting in the ‘continental’ geochemical signatures widely observed in Mesozoic mantle-derived mafic rocks from the North China Craton. The ‘continental’ geochemical signatures of Mesozoic mantle-derived mafic rocks suggest that lithospheric delamination could have occurred by the time of volcanic eruption in the northern margin of the North China Craton in the mid-Jurassic and later in the southern margin and Dabie–Sulu Orogen in the early Cretaceous.

The lithospheric thinning of the North China Craton during the Mesozoic has attracted considerable attention over the last two decades (e.g. Griffin et al. 1998; Guo F. et al. 2001; Zhang et al. 2002, 2003, 2004; Chen B. et al. 2003; Xu Y. et al. 2004a; Zhang 2005). Diamond-bearing kimberlites and mantle xenoliths demonstrate that a thick (c. 200 km) cold (c. 40 mW m22) lithosphere existed in the North China Craton in the Palaeozoic, but a thin (c. 80 km) and hot (c. 60 mW m22) lithosphere was present in the Cenozoic in the eastern part of the North China Craton (Eastern Block in Fig. 1; Griffin et al. 1998; Zheng et al. 2003). This indicates that about 120 km of lithosphere has been removed since the early Palaeozoic.

Also, the geochemical characteristics of the Palaeozoic and Cenozoic lithosphere mantle are very different (Table 1). The Palaeozoic lithospheric mantle underneath the North China Craton is characterized by EMII features, such as high 206 Pb/204Pb (c. 20.2), a significant variation of 87 Sr/86Sr, and negative 1Nd (c. 25) (Zheng & Lu 1999; Zhang et al. 2002), distinct from the Cenozoic lithospheric mantle below the Eastern Block of the North China Craton, which has Sr –Nd–Pb isotopic compositions similar to those of mid-ocean ridge basalt (MORB) and ocean island basalt (OIB) (Peng et al. 1986; Song et al. 1990; Basu et al. 1991). Apparently, the lithospheric mantle of the Eastern Block of the North China Craton was

From: ZHAI , M.-G., WINDLEY , B. F., KUSKY , T. M. & MENG , Q. R. (eds) Mesozoic Sub-Continental Lithospheric Thinning Under Eastern Asia. Geological Society, London, Special Publications, 280, 55–75. DOI: 10.1144/SP280.3 0305-8719/07/$15 # The Geological Society of London 2007.

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Fig. 1. Simplified map showing major locations of Mesozoic mantle-derived mafic rocks in Eastern China. Mesozoic lithospheric mantle provinces and locations of Mesozoic mantle-derived mafic rocks are modified from Li & Yang (2003) and Zhang et al. (2004). I, North margin (Liaoning region); II, Taihang; III, Luzhong; IV, Luxi; V, Jiaodong. Localities of Mesozoic mantle-derived mafic rocks from Eastern China are from Table 2. UHPM zone, ultrahigh-pressure metamorphic zone; WB, Western Block; TNCO, Trans-North China Orogen; EB, Eastern Block.

replaced between the Palaeozoic and Cenozoic (Zheng et al. 2003). The reason for the removal and replacement of the Palaeozoic lithospheric mantle is still not well known. Possible mechanisms include destabilization of the North China Craton as a result of the Indo-Eurasian collision (Menzies et al. 1993), replacement by asthenosphere upwelling (Xu Y. et al. 2004a), and destruction of the lithosphere as a result of the subduction of oceanic crust in the Palaeozoic and continental crust in the Mesozoic beneath both the northern and southern margins of the North China Craton (Zhang et al. 2003). Because the Mesozoic lithospheric mantle of the North China Craton is transitional, it can provide

critical constraints on understanding the lithospheric evolution during the Phanerozoic. Widespread Mesozoic mantle-derived mafic magmatism within the North China Craton provides important insights into the Mesozoic lithospheric mantle (Fig. 1 and Table 2). It is well known that Mesozoic mantlederived mafic igneous rocks are characterized by typical ‘continental’ geochemical signatures, including depletion of high field strength elements (HFSE), enrichment of large ion lithophile elements (LILE), negative 1Nd (most ranging from 210 to 220), variable 87Sr/86Sr (EMI-type with lower 87 Sr/86Sr and EMII-type with higher 87Sr/86Sr), and unradiogenic Pb isotope ratios (Qiu et al. 2000; Guo F. et al. 2001, 2003; Qiou et al. 2002;

EMII 0.704–0.711 25 19.2 –20.8

EMI 0.7049– 0.7078 24.9 to 29.1 16.77– 17.26

Liaoning EMI 0.7049 –0.7066 29.3 to 216.7 16.51 – 17.77

Taihang

Luzhong EMI 0.7040– 0.7069 24.0 to 221.1 17.04– 17.50

Mesozoic

EMII 0.7087 – 0.7114 29.8 to 218.0 16.57 – 17.65

Luxi – Jiaodong DMM 0.703– 0.706 22 to þ8 17.0– 18.6

Cenozoic lithospheric mantle

Data sources for mantle-derived rocks: Liaoning, Zhou et al. (2001), Zhang et al. (2003) and Yang et al. (2004b); Taihang, Chen B. et al. (2003, 2004), Zhang et al. (2004) and Wang et al. (2006); Luzhong, Guo F. et al. (2001), Xu Y. et al. (2004a) and Zhang et al. (2004); Luxi –Jiaodong, Qiou et al. (1997, 2002), Zhang et al. (2002), Xu Y. et al. (2004a), Yang et al. (2004a) and Ying et al. (2004). Palaeozoic lithospheric mantle is from Zheng & Lu (1999) and Zhang et al. (2002). Cenozoic lithospheric mantle is from Peng et al. (1986), Song et al. (1990) and Basu et al. (1991). DMM, depleted MORB mantle.

Mantle type 87 Sr/86Sr (130 Ma) 1Nd(130 Ma) 206 Pb/204Pb (130 Ma)

Palaeozoic lithospheric mantle

Table 1. Comparison of Sr – Nd– Pb isotopic compositions of Mesozoic mantle-derived mafic rocks and carbonatites from the North China Craton with Palaeozoic and Cenozoic lithospheric mantle

LOWER CRUST AND MESOZOIC MANTLE MAFIC ROCKS 57

130 – 145

Monzogabbro – diorite

Basalt

Linglong (16) Mouping (16)

Sulu orogenic belt Jimo (17)

K-rich basalts, lamproite Gabbro, diorite Alkaline basalt, olivine tholeiite Carbonatite

Gabbro Gabbro Gabbro, diorite

110 – 130

123.9– 132.5 120 + 1.1

106 – 107

118 – 122.9

115 – 124 126 – 132 124.9 + 1.8

115 120 + 5 126 – 132

155 125.2 + 4.5 178.3 + 3.8

135 – 145 150 – 160 155 129 – 138

122 – 127

Diorite, mafic enclave

Gabbro Gabbro Gabbro Monzogabbro, monzodiorite, syenite Gabbro Gabbro Xenolith-bearing volcanic pipe

Whole-rock K – Ar

142.4 + 2.2

Whole-rock K – Ar

Whole-rock K – Ar Whole-rock K – Ar

Phlogopite K – Ar, Rb – Sr isochron Whole-rock K – Ar

Whole-rock K – Ar SHRIMP zircon U – Pb Whole-rock K – Ar

Whole-rock K – Ar Strata Whole-rock K – Ar, SHRIMP zircon U – Pb

Zircon U – Pb, Rb – Sr isochron Whole-rock K – Ar Whole-rock K – Ar – Zircon U – Pb, Rb – Sr isochron – SHRIMP zircon U – Pb Whole-rock K – Ar

LA-ICP-MS zircon U – Pb

–

Dating method

–

Age (Ma)

Andesite, basaltic andesite, alkali basalt, sub-alkali basalt High-Mg andesite

Rock type

Olivine dolerite, gabbro, pyroxenite Basalt Basalt

Zibo basin

Laiwu –Zibo

Luxi –Jiaodong region Mengyin (9) Yinan (10) Fangcheng (10)

Luzhong region Jinan (6) Zouping (7) Laiwu (8)

Laiyuan (3) Hanxing Guyi and Fushan (4) Xishu, Wu’an, Hongshan (4) Donggang (4) Dongye (4) Xinyang (5)

Wulahada, Fuxin– Yixian (1) Liaodong (2) Taihang region Wang’an (3)

North China Block Liaoning region South of Chifeng – Kaiyuan fault, Liaoxi (1)*

Location

Table 2. Summary of Mesozoic mantle-derived rocks in Eastern China

Fan et al. (2001)

Yang et al. (2004a) Yang et al. (2004a)

Liu et al. (2004a)

Ying et al. (2004)

Qiou et al. (1997, 2002) Xu Y. et al. (2004) Zhang et al. (2002)

Lin et al. (1996) Guo F. et al. (2001) Xu Y. et al. (2004a); Zhang et al. (2004)

Zhang et al. (2004) Wang et al. (2006) Lu et al. (2003)

Zhang et al. (2004) Zhang et al. (2004) Zhang et al. (2004) Chen et al. (2004)

Chen et al. (2003a)

Yang et al. (2004b)

Zhang et al. (2003)

Zhou et al. (2001)

Reference

58 F. HUANG ET AL.

– Whole-rock K – Ar – Mineral K – Ar

130 98 – 105 – 164 + 2

Hornblende monzonite

–

120

Olivine basalts – shoshonite Tholeiite, alkaline basalt Gabbro

–

–

–

–

– –

Whole-rock K – Ar

SHRIMP zircon U – Pb, Rb– Sr and Sm – Nd isochron Rb – Sr, Sm – Nd isochron, Ar– Ar Zircon U – Pb

K – Ar

Whole-rock and mineral K–Ar SHRIMP zircon U – Pb, mineral Ar – Ar, zircon U–Pb Whole-rock and mineral K– Ar, SHRIMP zircon U–Pb

127.6– 131.8

130.2 + 1.4

120 – 133

122 – 128

116 – 132

*Numbers in parentheses indicate localities of Mesozoic mafic rocks from Eastern China shown in Figure 1.

Jiangxi (25) Jiangshan – Guangfeng (25) East– south of Guangxi (26) Mashan (27)

Middle – lower reach of the Yangtze River (23) SE Zhejiang (24)

Alkaline gabbro, trachyandesite, phonolite Gabbro, pyroxene diorite, basalt Basalt

Gabbro, pyroxenite, hornblendite Diabase, gabbro, lamprophyre, trachyandesite Gabbro

Zhujiapu (22)

Dongshichong, Dahuaping, Guanzhuang, Shutan, Anjiahe, Luoerling (22) Liujiawa (22) South China Block Ningwu (23)

Gabbro

Basaltic trachyandesite, trachyandesite Gabbro, diorite

108 – 124

201 – 215

Mafic enclave, pyroxene syenite, mafic dyke

Mafic dykes and enclave, biotite –pyroxene monzodiorite

122 – 127

Calc-alkaline lamproite

Jiaoziyan (22)

Shacun (22)

Dabie orogenic belt North Huaiyang (21)

Shijiusuo (20)

Sanjia lode-gold deposit (18) Jiazishan, Shidao (19)

Li et al. (2000)

Yu et al. (1993); Zhou et al. (1993); Yang et al. (1999) Liao et al. (1999) Yu et al. (2004) Guo X.-S. et al. (2001)

Yan et al. (2003)

Xing (1996)

Ma et al. (1998)

Wang et al. (2005)

Hacker et al. (1998); Jahn et al. (1999) Li et al. (1999)

Jahn et al. (1999); Zhao et al. (2005)

Fan et al. (2004)

Yang et al. (2005b)

Chen J. F. et al. (2003); Yang et al. (2005a)

Guo et al. (2004)

LOWER CRUST AND MESOZOIC MANTLE MAFIC ROCKS 59

60

F. HUANG ET AL.

Zhang & Sun 2002; Zhang et al. 2002, 2003, 2004; Chen B. et al. 2003; Li & Yang 2003; Xu Y. et al. 2004a, b; Yang et al. 2004a; Ying et al. 2004; Zhang et al. 2005). These signatures are not consistent with either the Palaeozoic or Cenozoic lithospheric mantle (Peng et al. 1986; Song et al. 1990; Basu et al. 1991; Chung 1999; Zhang et al. 2002; Table 1). It is widely considered that such ‘continental’ geochemical signatures of Mesozoic mantlederived mafic rocks from the North China Craton were derived from an enriched subcontinental lithospheric mantle (e.g. Guo et al. 2003; Yang et al. 2004a); two main models have been proposed. The first is that the subcontinental lithospheric mantle is an EMI-type resulting from multiple metasomatism related to subduction-related processes in the Archaean and Mesoproterozoic in the course of accretion of the North China Craton (e.g. Yang et al. 2004a); partial melting of the ancient subcontinental lithospheric mantle at different depths can explain the variation of geochemical features of Mesozoic mantle-derived mafic magmas (Guo et al. 2003). However, there is no evidence for the existence of an EMI-like enriched subcontinental lithospheric mantle with low 1Nd (up to 221) before the Mesozoic because the Palaeozoic kimberlites and peridotites have EMII-type isotopic features with a limited range of 1Nd from 25 to 27, high 206Pb/204Pb, and a large variation of 87Sr/86Sr from 0.705 to 0.712 (Zheng & Lu 1999; Zhang et al. 2002). The second model suggests that the Mesozoic subcontinental lithospheric mantle of the North China Craton was severely modified by a Si– Al-rich melt by partial melting of deeply subducted materials from the South China Block or PalaeoPacific plate (Zhang et al. 2002, 2003; Chen & Zhou 2005). This model is supported by mantle– melt reactions observed in olivine xenocrysts from Fangcheng basalts (Zhang 2005) and from composite dunite–orthopyroxene xenoliths captured in Laiwu high-Mg diorites (Chen & Zhou 2005). This model can explain the EMII-type signatures of mafic rocks and carbonatites from Western Shandong (Luxi) and Jiaodong peninsula (the Luxi– Jiaodong region hereafter), but it cannot explain EMI-type signatures in Mesozoic mafic rocks from the centre of the North China Craton (Taihang and Luzhong regions) where the effect of the subducted South China Block or PalaeoPacific plate is insignificant. Thus, the origin of the enriched signatures in Mesozoic mantle-derived mafic rocks remains controversial. This paper compiles recently published geochemical data from Mesozoic mantle-derived mafic igneous rocks to constrain the origin of their enriched signatures and understand the

transformation of the subcontinental lithospheric mantle of the North China Craton from the Palaeozoic to Cenozoic. The purpose of this study is (1) to reveal that geochemical signatures of Mesozoic mantle-derived mafic igneous rocks are consistent with the contribution of lower crust to the Mesozoic uppermost mantle of the North China Craton, but do not support derivation from a lithospheric mantle enriched by ancient fluid or melt metasomatism, and (2) to provide a delamination model to explain the lithospheric thinning process of the North China Craton during the Mesozoic.

Geological background Archaean rocks with ages of 3.6–3.8 Ga occur in the north to centre of the North China Craton, indicating that this is one of the oldest cratons in the world (e.g. Zheng et al. 2004, and references therein). However, the North China Craton is different from other old cratons in many respects, including high heat flow, thinned lithosphere, presence of earthquakes, unusually evolved bulk chemical crustal composition, and widespread magmatism from the late Mesozoic to Cenozoic (Gao et al. 2004). The North China Craton can be divided into the Western Block and Eastern Block, which are separated by the Trans-North China Orogen (Fig. 1). The North China Craton was stabilized when the Western and Eastern blocks collided along the Trans-North China Orogen at 1.8 Ga (Zhao et al. 2000). The North China Craton collided with the South China Block to the south along the Qinling –Dabie –Sulu Orogen in the early Triassic (Li et al. 1993) and to the north with the Centre Asian Orogen at the Solonker suture in the endPermian (Xiao et al. 2003). The North China Craton and attached southern Central Asian Orogen collided with northern Central Asian Orogen along the Mongol–Okhotsk suture in the Jurassic (Tomurtugoo et al. 2005). The Western Block of the North China Craton did not undergo lithospheric thinning or experience significant magmatism after the stabilization of the North China Craton at 1.8 Ga (Zhao et al. 2000; Zhang et al. 2003). However, magmatism occurred widely in the Eastern Block of the North China Craton after the Palaeozoic. The presence of Ordovician diamond-bearing kimberlites in the Eastern Block (e.g. in Mengyin and Fuxian) indicates that the lithosphere was cold and thick at that time (e.g. Griffin et al. 1998). Mesozoic magmatic rocks ranging from basalts to andesites and granites are widespread in Eastern China. Figure 1 and Table 2 show localities, rock types and ages of mafic intrusions in Eastern China. Mesozoic igneous carbonatites occur in the Luxi region

LOWER CRUST AND MESOZOIC MANTLE MAFIC ROCKS

(Ying et al. 2004). Mesozoic mantle-derived mafic rocks show distinct regional heterogeneity (Zhang et al. 2004). On the basis of geochemical differences of Mesozoic mantle-derived mafic rocks, the North China Craton is divided into five major units following a slightly modified scheme proposed by Zhang et al. (2004) (Fig. 1): Liaoning (province I); Tainang (province II); Luzhong (province III); Luxi (province VI); Jiaodong (province V). The geochemical features of these provinces are discussed below.

Geochemical database of Mesozoic mantle-derived mafic rocks from Eastern China The main purpose of this study is to understand the evolution of the lithospheric mantle of the North China Craton from the Palaeozoic to Cenozoic. Therefore, to avoid crustal-derived rocks, we have selected only basalts and basaltic andesites with a SiO2 content ,56 wt% and MgO content .4 wt% (most .5 wt%) from the data pool of Mesozoic magmatic rocks; the data sources are listed in Table 1. Some samples are mafic enclaves from Mesozoic granitoids, which Yang et al. (2004a, 2005a, b) suggested have mantle characteristics. Igneous carbonatites are from the Luxi region (Western Shandong) (Ying et al. 2004). For the northern margin of the North China Craton only samples older than 110 Ma were used because basalts later than 110 Ma in this area were produced by partial melting of asthenosphere in an extensional tectonic environment in the continental margin of Eastern Asia (Zhang et al. 2004). Coeval mantlederived mafic rocks from the South China Block and Dabie–Sulu Orogen were also studied for comparison. Trace element contents of most samples were analysed by inductively coupled plasma-mass spectrometry (ICP-MS) with precision better than +5% shown by more than two rock standards (e.g. Zhang et al. 2002; Yang et al. 2005a, b) or reproducibility of duplicated analyses (e.g. Guo F. et al. 2001). A few samples were measured by isotope dilution (ID) methods and X-ray fluorescence (XRF) spectrometry, including those of Jahn et al. (1999), which were useful to check for the consistency of Rb –Sr contents between the XRF and ID methods as a means of demonstrating the quality of trace element compositions. Because Sr –Nd–Pb isotopic compositions of the mantle-derived mafic rocks were measured in different laboratories, we corrected the isotopic ratios based on the same standard values: NBS-987 87Sr/86Sr ¼ 0.71025; BCR-1 143 Nd/144Nd ¼ 0.512630; NBS981 207Pb/204Pb ¼ 0.9142 + 0.0015. All initial isotopic ratios were calculated to 130 Ma.

61

Geochemistry of Mesozoic mantle-derived mafic rocks from the North China Craton Mesozoic mantle-derived mafic rocks share some common trace element compositions. Their trace element patterns are similar to those of continental crust (Rudnick & Gao 2003). They are enriched in light rare earth elements (LREE) relative to heavy rare earth elements (HREE) (Fig. 2a). Furthermore, mafic rocks from the North China Craton are enriched in LILE (Cs, Ba, U) and Pb, and depleted in HFSE (Nb, Zr, Ti) relative to N-MORB and OIB (Fig. 2b). Rb is also depleted relative to Ba, distinct from the upper crust but similar to the lower crust (Fig. 2b).

(a)

MC UC

OIB

N-MORB Mafic rocks from NCC

LC

MC UC

(b) Mafic rocks from NCC OIB

N-MORB

LC

Fig. 2. Chondrite-normalized REE patterns (a) and primitive mantle-normalized trace element spidergrams (b) of Mesozoic mafic rocks from the North China Craton (NCC). Chondrite and primitive mantle values are from Sun & McDonough (1989). Data sources of Mesozoic mafic rocks from the North China Craton are listed in Table 1. Data for Mesozoic mafic rocks from the South China Block are from Yu et al. (1993), Zhou et al. (1993), Xing (1996), Liao et al. (1999), Yang et al. (1999), Li et al. (2000), Guo X-S. et al. (2001), Yan et al. (2003) and Yu et al. (2004). OIB and MORB are from Sun & McDonough (1989); upper crust (UC), middle crust (MC), and lower crust (LC) are from Rudnick & Gao (2003).

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F. HUANG ET AL.

In the 1Nd (130 Ma)– 87Sr/86Sr (130 Ma) diagram (Fig. 3), Mesozoic mantle-derived mafic rocks and carbonatites from the North China Craton show regional variations and are distinct from mafic rocks in the South China Block (Zhang & Sun 2002; Zhang et al. 2002, 2003, 2004, 2005; Guo et al. 2003, 2004; Xu Y. et al. 2004a; Yang et al. 2004a, 2005a; Ying et al. 2004; Wang et al. 2006). Sr–Nd–Pb isotopic ratios of mafic rocks from the North China Craton are summarized in Table 1. 87 Sr/86Sr (130 Ma) increases gradually from the Taihang (province II) and Luzhong regions (province III) (87Sr/86Sr (130 Ma) ¼ 0.705–0.708) to the eastern North China Craton (Luxi–Jiaodong region, provinces IV and V with 87Sr/86Sr (130 Ma) ¼ 0.709–0.711), and the ranges of 1Nd (130 Ma) values are similar. The 1Nd (130 Ma) values of samples from the Luzhong range from 24.0 to 221 (Guo F. et al. 2001, 2003; Xu Y. et al. 2004a), Taihang from 29.3 to 216.7 (Chen B. et al. 2003; Zhang et al. 2004; Wang et al. 2006),

and Luxi–Jiaodong from 29.8 to 217.8 (Qiou et al. 1997, 2002; Xu Y. et al. 2004a; Yang et al. 2004a; Ying et al. 2004). Accordingly, the mantlederived mafic rocks from the North China Craton are divisible into two groups based on 87Sr/86Sr (130 Ma): the Liaoning (province I), Taihang (province II) and Luzhong regions (province III), which are characterized by EMI-like isotopic features (Lustrino & Dallai 2003), and the Luxi–Jiaodong region (provinces IV and V), which is characterized by its EMII-like isotopic character. Zhang et al. (2004) interpreted these features as evidence for the existence of a highly heterogeneous lithospheric mantle underneath Eastern China in the Mesozoic. The highest 1Nd (130 Ma) values of Mesozoic mantlederived mafic rocks are roughly in agreement with those of Palaeozoic kimberlites and peridotites, which have 1Nd (130 Ma) c. 25 (Zheng & Lu 1999; Zhang et al. 2002). Mesozoic mantle-derived mafic rocks from the North China Craton have Pb isotopic ratios with

10

MORB

SCB

0

Palaeozoic kimberlites and peridotites

εNd(130 Ma)

Liaoning SCB

–10 Taihang –20

Luxi–Jiaodong Dabie–Sulu

Luzhong Upper crust NCC mafic lower crust

–30

–40 0.702

0.705

0.708 87Sr/86Sr(130

0.711

0.714

0.717

Ma)

Fig. 3. Comparison of Sr –Nd isotopic ratios at 130 Ma of Mesozoic mantle-derived mafic samples from Eastern China. Data sources: North China Craton, Table 1; South China Block, Table 2; Dabie–Sulu Orogen, Ma et al. (1998), Jahn et al. (1999), Yang et al. (2004a, 2005a) Wang et al. (2005), Zhao et al. (2005) and Huang et al. (2007); MORB, Sun & McDonough (1989); Palaeozoic kimberlites and peridotites, Zheng & Lu (1999) and Zhang et al. (2002); lower crust, Zhang et al. (1998) and Liu et al. (2004b); upper crust, Jahn et al. (1999).

LOWER CRUST AND MESOZOIC MANTLE MAFIC ROCKS

significant regional variations (Fig. 4). Their Pb isotopic ratios recalculated at 130 Ma are lower than those of contemporary mafic rocks from the South China Block, and they are located to the left side of the 4.55 Ga geochron. Samples from the Luxi– Jiaodong region have higher 207Pb/204Pb (130 Ma) ratios than those from the Liaoning, Luzhong and Taihang regions. Compared with Mesozoic mantlederived mafic rocks from the Dabie Orogen (Wang et al. 2005; Huang et al. 2007), the mantle-derived mafic rocks from the North China Craton have similar uranogenic Pb isotopes (206Pb/204Pb and 207 Pb/204Pb) but lower thorogenic Pb isotopes 208 ( Pb/204Pb) (Fig. 4). The Palaeozoic kimberlites and peridotites from the North China Craton are characterized by high radiogenic Pb isotopic ratios (Fig. 4), which suggest that the contribution of Pb from the Palaeozoic enriched lithospheric mantle to the sources of Mesozoic mantle-derived mafic rocks in the North China Craton is insignificant. The geochemical signatures of the mafic igneous rocks could reflect the characteristics of the mantle source or result from crustal contamination by assimilation and fractional crystallization (AFC) during the magma ascent. Many workers have argued against significant crustal contamination based on lack of correlation between the Sr–Nd – Pb isotopic ratios and other geochemical features sensitive to the assimilation and fractional crystallization process (such as SiO2 content and Mg-number) (e.g. Zhang et al. 2004; Wang et al. 2005, 2006; Yang et al. 2005a; Zhao et al. 2005). However, with a limited range of SiO2 in this study (46–56 wt%), the AFC process might not be shown clearly by the correlation between SiO2 and Sr– Nd–Pb isotopes. Instead, Nb/U of Mesozoic mafic rocks from the North China Craton shows a large variation from 4.4 to 19, providing critical information on crustal contamination. The Nb/U of the bulk continental crust is c. 6.2 (Rudnick & Gao 2003), much lower than that of N-MORB and OIB (c. 47) (Hofmann et al. 1986). Thus crustal contamination during magma ascent can decrease Nb/U and change the Sr–Nd –Pb isotopic ratios of an evolved magma simultaneously, but fractional crystallization alone cannot change Nb/U and Sr–Nd –Pb isotopic ratios. The Nb/U ratio of Mesozoic mantle-derived mafic rocks for each individual province shows no obvious relationship with Sr –Nd–Pb isotopic ratios between the mantle and crust end-members in Figure 5. This precludes significant crustal contamination during the magma transport. However, it does not preclude the source mixing of three or more components. Samples from the Luxi– Jiaodong region have higher 87Sr/86Sr and lower Nb/ U than those from the Liaoning, Taihang and Luzhong regions (Fig. 5a). This may suggest that

63

a greater contribution of upper crustal material to the mantle source of the Luxi –Jiaodong samples compared with samples from other regions.

The possible contribution of lower crust to the subcontinental lithospheric mantle of the North China Craton Enrichment of LREE relative to HREE, high LILE/ HFSE, and a positive Pb anomaly are widely observed in arc magmas (e.g. Regelous et al. 1997). Arc magmas are produced as a result of partial melting of the overlying mantle wedge metasomatized by slab-derived fluids derived from subducted oceanic crust and sediments with high LREE/HREE and LILE/HFSE ratios as well as a positive Pb anomaly (Brenan et al. 1995; Keppler 1996; Kogiso et al. 1997; Peate et al. 2001; Manning 2004). Such fluids also have EMII-type radiogenic Sr and Pb isotopic ratios (e.g. Regelous et al. 1997). Accordingly, mantle-derived mafic rocks form the Luxi –Jiaodong region (provinces IV and V) could be derived from enriched lithospheric mantle metasomatized by fluid derived from ancient continental sediments during subduction-related processes (Zhang et al. 2002) based on their high 87Sr/86Sr (130 Ma) and 207 Pb/204Pb (130 Ma) as well as low Nb/U ratio. However, the subduction-related, fluid-addition model cannot explain the EMI-type Sr–Nd– Pb isotopic ratios in Mesozoic mantle-derived mafic rocks from the Liaoning, Taihang and Luzhong regions. Alternatively, the EMI-type isotopic signatures of Mesozoic mantle-derived mafic rocks could reflect the characteristics of the subcontinental lithospheric mantle of the North China Craton, caused by former metasomatism that formed phlogopite-bearing lithospheric mantle (Guo et al. 2003; Yang et al. 2004a). The EMI-type subcontinental lithospheric mantle with its extremely low 1Nd(t) values and unradiogenic Sr and Pb isotopes, similar to that which gave rise to the Smoky Butte lamproites (Fraser et al. 1985), is a possible source. However, the lower Rb/Ba and much higher Ce/Pb ratios of the Smoky Butte lamproites than the mantle-derived mafic rocks in the North China Craton argue against a major contribution from an EMI-like subcontinental lithospheric mantle to Mesozoic mantle-derived mafic rocks from the North China Craton (Fig. 6a). As shown in Figure 6, the high Ba/Nb and La/Nb, and low Ce/Pb and Nb/U ratios of Mesozoic mantlederived mafic rocks from the North China Craton share for more affinities with the crustal estimates, but are clearly different from those of the OIB, N-MORB (Hofmann et al. 1986), and primitive mantle (Sun & McDonough 1989). EMI-type

64

F. HUANG ET AL.

(a)

Geo chr on

16.0

207Pb/ 204Pb(130

Ma)

15.8

Palaeozoic kimberlites and peridotites

EMII RL

SCB

15.6

NH

EMI 15.4

Da 15.2

15.0

bie

Lower crust 15

16

17

18

19

206Pb/ 204Pb(130

40

(b)

Ma)

21

Ma)

EMII SCB

208Pb/ 204Pb(130

20

38

Palaeozoic kimberlites and peridotites

N-MORB

EMI ie

b Da

Lower crust

36

L

R NH

34 15

16

17

18

206Pb/ 204Pb(130

19

20

21

Ma)

Fig. 4. Pb isotopic compositions calculated at 130 Ma of Mesozoic mantle-derived mafic rocks from Eastern China. Data source: Dabie, Wang et al. (2005) and Huang et al. (2007); Liaoning, Zhou et al. (2001) and Zhang et al. (2003); Taihang, Zhang et al. (2004) and Wang et al. (2006); Luzhong, Xu Y. et al. (2004a); Luxi– Jiaodong, Qiou et al. (1997, 2002), Zhang et al. (2002), Xu Y. et al. (2004b) and Yang et al. (2004a); South China Block (SCB), Chen et al. (1994), Zhang (1995) and Yan et al. (2003); MORB, Zindler & Hart (1986); EMI and EMII are from Lustrino & Dallai (2003); lower crust, Tu et al. (1993). NHRL, Northern Hemisphere Reference Line (Hart 1984). (207Pb/204Pb)NHRL ¼ 0.1084  (206Pb/204Pb) (130 Ma) þ 13.491; (208Pb/204Pb)NHRL ¼ 1.209  (206Pb/ 204Pb) (130 Ma) þ 15.627.

LOWER CRUST AND MESOZOIC MANTLE MAFIC ROCKS 0.714

(a)

Upper crust

(a) UC

N-MORB & OIB

Rb/Ba

0.708 Lower crust

LC Dabie-Sulu

0.01

N-MORB

0.705

Liaoning Luxi-Jiaodong Taihang Luzhong

1E-3

0.702

W.A. Lamproite

0.1

87

Sr/86Sri

0.711

1

Liaoning Taihang Luzhong Luxi-Jiaodong

0

10

20

10

30 Nb/U

40

10 Ce/Pb

1

1000

100

(b)

(b) Ba/Nb

N-MORB

0 εNd(t)

Smoky Butte Lamproite

50

Dabie-Sulu

100

UC

LC Primitive mantle

10

OIB

–10 1

N-MORB

1

–20

Lower crust

Upper crust

0 19

10 Upper crust

20

30 Nb/U

30

50

La/Nb

10

(c) N-MORB & OIB

20

N-MORB Ce/Pb

(c)

40

18

Dabie-Sulu

10 LC

17

UC

0

206

Pb/ 204Pbi

65

16

0

10

20

30 Nb/U

40

50

Lower crust

15

0

10

20 30 Nb/U

40

50

Fig. 5. Correlation of Nb/U with Sr –Nd –Pb isotopic compositions of Mesozoic mantle-derived mafic rocks from the North China Craton. Nb/U ratios of N-MORB and upper and lower crust are from Hofmann et al. (1986) and Rudnick & Gao (2003), respectively. Sr–Nd–Pb isotopic compositions: lower crust, granulite xenoliths from Zhang et al. (1998) and Liu et al. (2004b); upper crust, Xu Y. et al. (2004b); mafic rocks from the North China Craton, Table 1.

lithospheric mantle recently discovered in Taihang region has 1Nd (130 Ma) ranging from 26.9 to 210.6 with Ce/Pb ratios from 41.5 to 72.0 (Ma & Xu 2006), higher than most Mesozoic mantlederived mafic rocks from the North China Craton. This reinforces the involvement of crustal materials in the source of Mesozoic mantle-derived mafic rocks from the North China Craton.

Fig. 6. Variation of trace element ratios of Mesozoic mantle-derived mafic rocks from the North China Craton. The crustal values of the upper and lower crust are from Rudnick & Gao (2003). Data sources for other fields: Western Australia (W.A.) and Smoky Butte lamproites, Fraser et al. (1985); OIB and MORB, Hofmann et al. (1986). The high Ba/Nb and La/Nb, and low Nb/U and Ce/Pb ratios of Mesozoic mafic rocks from Eastern China relative to the mantle values show clear involvement of crustal materials.

Moreover, fluid-related metasomatism can increase Rb/Sr, Pb/U and Nd/Sm ratios of the lithospheric mantle, which will generate high 87Sr/86Sr, and low 206Pb/204Pb and 143Nd/144Nd with time, and a negative correlation between the Sr and Pb isotopic ratios (Hawkesworth et al. 1990a, b). This is supported by the negative correlation between the Sr and Pb isotopic ratios of the peridotite xenoliths

66

F. HUANG ET AL.

correlation with 206Pb/204Pb (130 Ma) as well as the samples from the Taihang and Liaoning regions, which is not consistent with the possible predicted scenario of an ancient fluid metasomatism. Wang et al. (2006) suggested that the low 87Sr/86Sr

from Kimberley, South Africa and that of lamproites from Western Australia (Fraser et al. 1985; Hawkesworth et al. 1990a, b). However, as Figure 7 shows, the 87Sr/86Sr (130 Ma) values of the samples from the Luxi–Jiaodong region show a slightly positive

(a)

(130 Ma)

50

20

30

10

87Sr/ 86Sr

10

6 4

20 10

6

2 4

2

1 1

206Pb/ 204Pb

(130 Ma) (b)

2

1

1 2

εNd (130 Ma)

4 6

4 6

10

10 20 30 50 10

206Pb/ 204Pb

30

50

(130 Ma)

Fig. 7. Sr– Nd– Pb isotopic composition of Mesozoic mantle-derived mafic rocks from the North China Craton. The data sources are as in Figures 3 and 4. The results of a source-mixing model between continental crust and depleted MORB mantle (DMM) show that Mesozoic mantle-derived mafic rocks from the North China Craton can be interpreted as derived from DMM modified by lower crust with a few percent of upper crust. Proportions of components are marked in wt%. The parameters used in the mixing calculation are listed in Table 3. PKP, Palaeozoic kimberlites and peridotites from the North China Craton (Zheng & Lu 1999; Zhang et al. 2002).

LOWER CRUST AND MESOZOIC MANTLE MAFIC ROCKS

(130 Ma), 1Nd (130 Ma) and Pb isotopic ratios observed in basaltic rocks from the Taihang region might result from an ancient metasomatism caused by an SiO2-rich melt derived from subducting plate during the collision between the Western and Eastern Blocks of the North China Craton. However, high polymerization of the SiO2-rich melt can enhance the partition coefficient of Sr more than Rb because Sr2þ has a larger charge/radius radio than Rbþ (Ryerson & Hess 1978; Huang et al. 2006). Such a SiO2-rich melt could also have high Rb/Sr. For instance, Wulff-Pedersen et al. (1999) reported that SiO2-rich glasses in mantle xenoliths (sample PAT2-4, PAT2-68 and PAT2-41) have a high Rb content (up to 274 ppm) and Rb/Sr (up to 0.6). The metasomatized lithospheric mantle will produce a high 87Sr/86Sr with time, not consistent with the observations on the Taihang samples. Actually, the positive correlation between 87Sr/86Sr (130 Ma) and 206Pb/204Pb (130 Ma) is a typical feature of continental crustal rocks. Ancient lower crustal rocks have lower 87Sr/86Sr and much lower 206 Pb/204Pb, whereas upper crustal rocks have higher 87Sr/86Sr and 206Pb/204Pb. Therefore, the positive correlation between 87Sr/86Sr (130 Ma) and 206 Pb/204Pb (130 Ma) of the mantle-derived mafic rocks may suggest the involvement of continent crustal materials in their source (Lustrino et al. 2007). A source-mixing modelling reveals that the contribution of lower continental crust (or with a few percent of upper continental crust) to the depleted MORB mantle can produce the Sr–Nd–Pb isotopic features of Mesozoic mantle-derived mafic rocks in the North China Craton (Fig. 7, Table 3). The samples from the Luxi–Jiaodong region (provinces IV and V) require a higher proportion of upper crust in their mantle source than those from the centre of the North China Craton far away from the Phanerozoic subduction zones, which might be due to the metasomatism of SiO2-rich melts related to the subduction of the South China Block (Zhang et al. 2002; Zhang 2005) or Palaeo-Pacific Ocean to the North China Craton (Chen & Zhou 2005). In summary, Mesozoic mafic rocks from the North China Craton cannot simply result from

67

partial melting of the subcontinental lithospheric mantle enriched by a subduction-related fluid or ancient fluid or melt metasomatism, but are due to involvement of continent crustal materials (mostly lower crust). Notably, the contribution of the lower continental crust has been recognized in the genesis of Plio-Pleistocene tholeiitic and alkaline volcanic rocks in Sardinia (Italy) (Lustrino et al. 2000, 2007). In this case, low Nb/ U, Ce/Pb and 206Pb/204Pb values have been used as evidence for involvement of the lower crust in the volcanic rocks that have low radiogenic Pb isotopic ratios.

How was the lower crust incorporated into the uppermost mantle? Lower continental crust can be recycled and modify geochemical features of the upper mantle in subduction and continental collision zones as a result of deep subduction (Huang et al. 2007) or lithospheric delamination (e.g. England 1993; Kay & Kay 1993; Lee et al. 2000; Gao et al. 2004; Lustrino 2005). Because the North China Craton has been stable for 1.8 Ga (Zhao et al. 2000), and the subducted South China Block has different Pb isotopic ratios compared with the lower crust and Mesozoic lithospheric mantle of the North China Craton (Huang et al. 2007), we propose that the lower crust of the North China Craton was incorporated into the upper mantle by lithospheric delamination. Briefly, underneath over-thickened lithosphere caused by oceanic subduction or continental collision, high-pressure metamorphism can lead to the formation of eclogite or garnet pyroxenite in the lower continental crust (e.g. Kay & Kay 1993; Gao et al. 2004). The density of the garnet-bearing metamorphic rocks can be as high as 3.8 g cm23 depending on the quantity of garnet, which has a density higher than that of lithospheric and asthenospheric mantle (c. 3.3 g cm23) (Lustrino 2005, and references therein). Therefore, eclogitic lower crust and lithospheric mantle might sink into warmer mantle because of its negative buoyancy.

Table 3. Parameters for source mixing between lower crust and mantle components 87

Sr/86Sr

*

DMM Lower crust† Upper crust†

0.703 0.709 0.718

Sr (ppm)

1Nd

Nd (ppm)

20 348 320

8 233 225

1.2 11 27

206

Pb/204Pb 18 16.2 20

Pb (ppm) 0.2 4 17

*Sr –Nd data for the depleted MORB mantle (DMM) are from Jahn et al. (1999); Pb data are from Sun & McDonough (1989). † Sr –Nd –Pb contents of the lower and upper crust from Rudnick & Gao (2003). Isotope data: lower crust, lower crustal xenoliths from Liu et al. (2004b); upper crust from Xu Y. et al. (2004b).

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Many geological data from of the North China Craton are remarkably consistent with the proposed scenarios of the lithospheric delamination model. 1. An over-thickened lithosphere could have been present in the Mesozoic as a result of the collision of the North China Craton with the South China Block in the south and with the Central Asian Orogenic Belt in the north in the Permian– Triassic (Li et al. 1993; Zhang et al. 2003). Discovery of eclogite xenoliths in Mesozoic diorite intrusions in the Xu–Su region with an inherited zircon U –Pb age of 2.4–2.5 Ga and a metamorphic zircon U– Pb age of c. 206 + 15 Ma as well as a Sm–Nd age of 219.4 Ma indicates that an over-thickened mafic lower crust with eclogite facies existed in the North China Craton in the late Triassic (Gao et al. 2004; Xu Y. et al. 2004). 2. Petrological and geochemical evidence from Late Jurassic high-magnesium andesites, dacites and adakites in the North China Craton demonstrates that foundering of mafic lower continental crust into convecting upper mantle occurred in the North China Craton (Gao et al. 2004). 3. Granulite xenoliths entrained in the Hannuoba basalts indicate that the Precambrian lower crust of the North China Craton has unradiogenic Pb, low 1Nd(t), and variable radiogenic Sr (Zhang et al. 1998; Liu et al. 2004b), in agreement with the isotopic signatures of Mesozoic mantle-derived mafic rocks from the North China Craton. 4. Study of Mesozoic lower crustal xenoliths also indicates that up to 10 km of Mesozoic mostly lower crust was delaminated into the upper mantle before Cenozoic basaltic magmatism (Zheng et al. 2003). 5. Widespread coeval granitic rocks coexist with the mantle-derived rocks in the North China Craton, which was a thermal high corresponding to the lithospheric delamination (Wu et al. 2003a, b). Following lithospheric delamination, a hot asthenospheric mantle rises to replace the volume previously occupied by detached mafic lower crust and lithospheric mantle. Consequently, decompression melting of the rising asthenospheric mantle took place, creating a basaltic melt with a depleted isotopic signature. Thus, the geochemical signature of the mantle-derived mafic rocks should change abruptly shortly after the delamination. Therefore, the temporal variation of the geochemical signatures of the mantle-derived mafic rocks could have provided the critical temporal constraint on the lithospheric thinning (e.g. Xu Y. et al. 2004a). However, the abrupt variation in geochemical features of the mantle-derived mafic magma from the North China Craton developed more than 40 Ma after the lithospheric delamination. For instance, volcanic eruptions occurred several times in Western Liaoning from the mid-Jurassic to Cenozoic,

providing a good opportunity to test the geochemical correspondence to the lithospheric delamination and the mantle-derived mafic magmas (Zhang et al. 2003; Gao et al. 2004). Gao et al. (2004) suggested that the delamination was initiated at 159 Ma. As Figure 8 shows, there is no significant variation in 1Nd(t) of Mesozoic basaltic rocks from Western Liaoning from 166 Ma (the Lanqi Formation) to 125 Ma (the Yixian Formation) until the Zhanglaogongtun (ZLGT) Formation at 106–90 Ma. Thus the 1Nd(t) values of the mantle-derived mafic magma were constant for a long period after the start of delamination. This is not in agreement with the abrupt geochemical variation as suggested in the previous delamination model. Therefore, although the delamination model can explain the aspects mentioned above, there is still a contradiction between the model and the data. The new delamination model of Lustrino (2005) might solve the contradiction by assuming that the asthenosphere was modified by delaminated lower crust before partial melting. Delaminated mafic lower crust can undergo partially melting to produce a tonalite–trondhjemite–granodiorite magma with a crustal geochemical signature. SiO2-rich melts tend to percolate upwards and erupt as adakitic magma (e.g. Xu et al. 2002; Gao et al. 2004) or they metasomatize the uprising asthenospheric mantle, leading to continental geochemical characteristics (Lustrino 2005). After the lithospheric delamination and detachment, new lithospheric mantle with a strong crustal signature forms by cooling of the asthenospheric mantle, which replaces the volume formerly occupied by the sunken lithospheric mantle and lower crust. Such metasomatized mantle may be reactivated by regional extension tectonics, producing a mantle-derived mafic magma with a similar ‘continental’ geochemical signature (Lustrino 2005). Alternatively, the ‘continental’ geochemical signatures of the Yixian Formation basalts (124 Ma) could be due to a second delamination occurring in the Early Cretaceous in the Liaoxi region. Because it is highly unlikely that the lithospheric delamination can occur in the same area several times, the distribution of the Lanqi and Yixian Formations should be spatially separated in different areas. However, this is not consistent with the geographical observation in the western Liaoning region, where the Lanqi and Yixian Formations are developed in the almost the same area (e.g. Beipiao).

A lithospheric thinning model As discussed above, lithospheric delamination played an important role in lithospheric thinning below the North China Craton. Here, we propose a

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10 Yang, W. unpublished data Zhang et al. 2004 Ji et al. 2004

ZLGT formation

5

εNd(t)

0

–5

–10

–15 180

Yixian formation

Lanqi formation

160

140

120

Magmatic hiatus

100

80

t (Ma) Fig. 8. Variation of 1Nd(t) with time of the mafic rocks from Western Liaoning. Data sources: Lanqi formation, W. Yang, unpubl. data; Yixian, Ji et al. (2004) and W. Yang, unpubl. data; ZLGT (Zhanglaogongtun) Formation, Zhang et al. (2004).

geodynamic model (Fig. 9) showing that Palaeozoic subduction zones around the North China Craton were critical to the over-thickening and later thinning of the lithosphere, and that interaction between the delaminated lower crust and upwelling asthenosphere was responsible for the ‘continental’ signature of Mesozoic mantle-derived mafic rocks. It has been suggested that the lithospheric thickening of the North China Craton was caused by continental collision between the North China Craton and South China Block (Gao et al. 2004; Zhang 2005), subduction of the Palaeo-Pacific slab (Tatsumoto et al. 1992; Wu et al. 2005), or collision of the North China Craton–Mongolia with the Siberian plate during closure of the Mongol – Okhotsk Ocean. Although the mechanism of the lithospheric thickening is still an open question, the subduction-related events surrounding the North China Craton from the late Palaeozoic to early Jurassic are important because they may be responsible for the condition of hydrous fluids to weaken the lithospheric mantle (Fig. 9a), resulting in an over-thickened lithosphere during the continental collisions (Fig. 9b) (B. F. Windley, talk in Symposium on Mesozoic Lithospheric Evolution of North China and Adjacent Regions, 2005). Metamorphic reactions occurred in the lower mafic continental crust with the formation of eclogite and garnet pyroxenite beneath the North China Craton leading to negative buoyancy in over-thickened

lithosphere in the early Mesozoic. The unstable lithosphere may then have delaminated from the overlying lithosphere above and subsided into the upper mantle (Lustrino 2005; Fig. 9c). Timing of the delamination event in the North China Craton is still controversial. Because the ages of most Mesozoic igneous rocks in the North China Craton cluster around 130 Ma (Wilde et al. 2003; Gao et al. 2004; Xu Y. et al. 2004a; Wu et al. 2005), Wu et al. (2005) suggested that the lithospheric delamination in Eastern China occurred in the early Cretaceous, resulting from Kula –Pacific plate subduction, possibly aided by a superplume associated with global-scale mantle upwelling. However, because a high-Mg adakite in Western Liaoning region has an age of 159 Ma, the delamination should have begun in the middle Jurassic (Gao et al. 2004). As discussed above, because the ‘continental’ geochemical signature of the mantle-derived mafic rocks from the North China Craton is due to the contribution of lower continental crust to the mantle source, the first magma event producing mafic rocks with these signatures should be a sharp response to lithospheric delamination. The negative 1Nd(t) of the Lanqi Formation basalts (166 Ma) indicates that the upper mantle beneath Western Liaoning had already been modified by delaminated lower crust. Thus lithospheric delamination should have occurred in Western Liaoning in the mid-Jurassic (Fig. 9c).

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F. HUANG ET AL. S SCB

(a) Palaeozoic (~460 Ma) NCC

N MB

Kimberlite Mongolian Ocean

Tethyan Ocean

id Flu

(b) Triassic to Early Jurassic (220–190 Ma) SCB

NCC

Dabie-Sulu orogen

MB

Mongo 1–0 khotsk Ocean

nic

ea

Oc cru st

(c) Mid-Jurassic (~160 Ma) SCB

NCC

Dabie-Sulu orogen

(d) Early–Cretaceous (~130 Ma) SCB

NCC

(e) Cenozoic (90 Ma to present) SCB

Dabie-Sulu orogen Basalts

NCC

Mongolo-0 khotsk

Fig. 9. Schematic illustration of the stages of evolution of the lithosphere of the North China Craton (NCC) from the Palaeozoic to Cenozoic. (a) Palaeozoic (c. 460 Ma): the Palaeo-Tethyan Ocean and Mongolian Ocean subducted towards the North China Craton from the south and north, respectively; hydrous fluids were released from the subducted oceanic slabs and metasomatized the overlying mantle wedge; kimberlite with peridotite xenoliths was developed in the North China Craton. (b) Triassic to Early Jurassic (220–190 Ma): collision of the North China Craton with South China Block (SCB) along the Dabie–Sulu Orogen and with the Mongolian Block (MB) along the Yanshan

LOWER CRUST AND MESOZOIC MANTLE MAFIC ROCKS

Partial melts from the delaminated lower crust metasomatized the upwelling asthenospheric mantle that replaced the volume formerly occupied by the sunken lithospheric mantle and the lower crust (Fig. 9c and d). Regional extension in the early Cretaceous led to partial melting of the new metasomatized mantle and crust in Western Liaoning, producing igneous rocks with variable chemical compositions from basalts to andesites and granites (Fig. 9d). Because no mid-Jurassic high-Mg adakitic rocks have yet been observed in the southern margin of the North China Craton and Dabie –Sulu Orogen, where all reported adakitic rocks formed in the early Cretaceous (Xu J. F. et al. 2002; Xu W. L. et al. 2006), the lithospheric delamination in the southern margin of the North China Craton and Dabie –Sulu Orogen probably occurred in the early Cretaceous (Fig. 9d). Thus the delamination of lithosphere of the North China Craton took place earlier in the northern margin than the southern margin and the Dabie – Sulu Orogen. It is possible that the delamination in the northern margin could be related to the closure of the Okhotsk Sea in the late Jurassic, which formed large-scale thrust faults in the Yanshan belt in Northern China (the northern margin) but had no significant effect on the southern margin (Davis et al. 1998). The delamination of the southern margin of the North China Carton and Dabie –Sulu Orogen could be triggered by oblique westward subduction of the Palaeo-Pacific plate at c. 130 Ma (Tatsumoto et al. 1992; Wu et al. 2005). Further studies focusing on the tectonic relationship between the North China Craton and its adjacent regions will help resolve the time variation of delamination. We emphasize that the entire lithospheric thinning in the North China Craton was unlikely to have been caused by any single event. Lithospheric delamination could have reduced the lithosphere thickness to a normal level (c. 110 km). In the late Cretaceous –Cenozoic, the lithosphere of the North China Craton was thinned further by

71

NNE –SSW extension in Eastern Asia, which resulted in partial melting of upwelling asthenosphere mantle with depleted isotopic signatures (Zhang et al. 2003) (Fig. 9e).

Conclusions Mesozoic mantle-derived mafic rocks from the North China Craton have geochemical signatures similar to those of ancient lower continental crust in trace element and Sr–Nd –Pb isotopic compositions. 87Sr/86Sr (130 Ma) of Mesozoic mantlederived mafic rocks from the North China Craton correlate positively with 206Pb/204Pb (130 Ma). Because an enriched mantle metasomatized by ancient fluids or melts produces a negative correlation between 87Sr/86Sr (130 Ma) and 206Pb/ 204 Pb (130 Ma), and because of the low Ce/Pb and Nb/U, and high Ba/Nb and La/Nb of Mesozoic mafic rocks from the North China Craton, the geochemical features mentioned above argue against derivation from an old enriched subcontinental lithospheric mantle. Instead, the typical ‘continental’ signatures in Mesozoic mantle-derived mafic rocks from the North China Craton mainly reflect the contribution of lower mafic continental crust to the uppermost mantle. Much geochemical and petrological evidence indicates that the continental lower crust and lithospheric mantle of the North China Craton could have been delaminated and sunk into the upper mantle in the Mesozoic. Melting of the delaminated continental lower crust would have created SiO2-rich melts, which metasomatized the upper mantle and resulted in the ‘continental’ geochemical signatures of Mesozoic upper mantle of the North China Craton. Thus, magma events, producing the mantle-derived mafic rocks with these ‘continental’ geochemical features, provide critical constraints on the timing of lithospheric delamination. According to the temporal variation of geochemical features of the Mesozoic mafic

Fig. 9. (Continued) belt resulted in an over-thickened lithosphere. (c) Mid-Jurassic (160 Ma): the Mongol–Okhotsk Ocean was closed and the Siberian plate collided with the North China Craton–Mongolia plate; the eclogitic mafic lower crust underneath the northern margin of the North China Craton was delaminated together with the lithospheric mantle as a result of negative buoyancy; SiO2-rich melts from the lower crust metasomatized the asthenospheric mantle, filling the vacancy of the delaminated lithosphere; volcanic rocks were developed in the Lanqi Formation (Lanqi FM). (d) Early Cretaceous (130 Ma): lithosphere underneath the southern margin of the North China Craton and Dabie Orogen was delaminated; partial melting of the metasomatized lithospheric mantle took place widely in an extensional tectonic environment, producing Mesozoic mantle-derived mafic rocks; crustal melting produced granitic intrusions; basaltic magmas were underplated beneath the crust; post-collisional magmatism also occurred in the Dabie–Sulu and Yanshan belts. (e) Cenozoic (c. 90 Ma to present): NNE– SSW extension along the continental margin of Eastern Asia thinned the lithosphere further; the lithospheric mantle of the North China Craton was completely replaced by newly accreted lithospheric mantle; Cenozoic basaltic magmas with depleted Sr– Nd– Pb isotopic ratios were developed in some extensional basins.

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rocks, the lithospheric thinning might have happened by the time of volcanic eruptions in Western Liaoning in the mid-Jurassic and in the southern margin of the North China Craton and Dabie –Sulu Orogen in the early Cretaceous. With further delamination and extension, the Palaeozoic lithosphere of the North China Craton was completely replaced by fertile lithospheric mantle in the Cenozoic. We thank C. Lundstrom for suggestions on an earlier version of our manuscript. M. G. Zhai, Q. R. Meng, B. Windley and J. H. Yang are thanked for their considerable contribution to this special volume. F. Dong provided some Chinese references. Comments by M. Lustrino and an anonymous reviewer improved our manuscript significantly. We are grateful to B. Windley for help with English polishing. This work was funded by the Natural Science Foundation of China (Grant 40573010).

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Late Mesozoic mafic magmatism from the North China Block: constraints on chemical and isotopic heterogeneity of the subcontinental lithospheric mantle W.-M. FAN1, F. GUO1, Y.-J. WANG1 & H.-F. ZHANG2 1

Key Laboratory of Marginal Sea Geology, Guangzhou Institute of Geochemistry, Chinese Academy of Sciences, Wushan, Guangzhou, 510640, China (e-mail: [email protected])

2

Institute of Geology and Geophysics, Chinese Academy of Sciences, Beijing, 100029, China Abstract: Available major, trace element and Sr– Nd isotope data for the late Mesozoic mafic rocks in the eastern North China Block (NCB) show chemical and isotopic differences between rocks from different tectonic units. Such differences are interpreted as signatures inherited from the melted mantle sources, which had experienced distinctive enrichment processes during lithospheric evolution. The subcontinental lithospheric mantle beneath the NCB interior is characterized by long-term light REE (LREE) enrichment and EM1-like Sr–Nd isotopic signatures. Such a lithospheric mantle is mainly composed of chemically refractory peridotites that are common in cratonic regions. In contrast to that of the NCB interior, beneath the northern part of the NCB a relatively chemically fertile mantle was enriched in large ion lithophile elements and LREE and depleted in Nb– Ta and Th–U. It has higher 87Sr/86Sr(i) and 1Nd(t) than that of the interior of the block, and is interpreted to have been modified by recycled lower continental crust components related to the palaeo-Asian Ocean subduction. The lithospheric mantle beneath the southern NCB has the highest 87Sr/86Sr(i) and the lowest 1Nd(t), and is chemically transitional between the interior and northern part of the block. Formation of such an enriched lithospheric mantle was closely associated with modification from the subducted Yangtze lower–middle crust during Triassic collision between the North China and Yangtze Blocks. A lithospheric extension–thinning model is proposed to explain the petrogenesis of these late Mesozoic mafic rocks in the eastern North China Block. This process was amplified by effects from surrounding plate interactions, including the rapid northward movement of the palaeo-Pacific Ocean, compressional forces from the Siberian plate, the Tethyan tectonic belt and possibly the Indo-China Block. The resultant forces triggered lithospheric extension, asthenospheric upwelling, and decompressional melting of the enriched mantle sources.

The nature of the subcontinental lithospheric mantle (SCLM) is important for understanding the Earth’s evolution and basalt petrogenesis. It is widely considered that certain kinds of mantle reservoirs correspond to regional tectonic settings (e.g. Hart 1988; Boyd 1989; Menzies 1989). For instance, beneath cratonic regions such as the Kaapvaal, Siberia, Wyoming and Slave Cratons (Walker et al. 1989; Carlson & Irving 1994; Pearson et al. 1995; Griffin et al. 1999), the lithospheric mantle is generally characterized by low Rb/Sr and Sm/Nd as well as an I-type enriched mantle (EM1)-like isotopic signature. A II-type enriched mantle (EM2)-like mantle reservoir is located beneath circum-cratonic terranes, where crustal material had been recycled into the mantle through plate subduction (Boyd 1989), whereas there are mid-ocean ridge basalt (MORB) and ocean island basalt (OIB) types of mantle reservoirs beneath tectonically mobile Phanerozoic belts (Menzies 1989). However, some tectonically remobilized regions, such as the eastern North China Block (NCB) underwent

decoupling between crustal rocks and subcrustal mantle. The NCB is one of the oldest continents, with recorded crustal ages of c. 3.8 Ga (Liu et al. 1992). In contrast, the present lithospheric mantle is mainly composed of chemically fertile spinel peridotite with a lithospheric thickness of 70 – 80 km, similar to that beneath oceanic plates or Phanerozoic mobile belts (e.g. Griffin et al. 1992, 1998; Menzies et al. 1993; Fan et al. 2000). Comparative studies of mantle xenoliths hosted in early Ordovician diamondiferous kimberlites and Cenozoic alkaline basalts suggest a loss of .100 km of lithospheric mantle during the last 400 Ma (Fan & Menzies 1992; Griffin et al. 1992, 1998; Menzies et al. 1993; Chi & Lu 1996; Wang et al. 1998; Fan et al. 2000; Xu 2001; Zheng et al. 2001). Such a thinning process has attracted the interest of international geologists, and several tectonic models have been proposed to explain the transformation of the early Palaeozoic subcratonic mantle lithosphere (180– 220 km, within the diamond stability field) into the Cenozoic

From: ZHAI , M.-G., WINDLEY , B. F., KUSKY , T. M. & MENG , Q. R. (eds) Mesozoic Sub-Continental Lithospheric Thinning Under Eastern Asia. Geological Society, London, Special Publications, 280, 77–100. DOI: 10.1144/SP280.4 0305-8719/07/$15 # The Geological Society of London 2007.

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MORB– OIB-type mantle (e.g. Fan & Menzies 1992; Menzies et al. 1993; Gao et al. 1998, 2004; Fan et al. 2000; Xu 2001; Zheng et al. 2001). As an intermediate stage of lithospheric evolution, the Mesozoic lithospheric mantle holds the key to the following questions: (1) What kind of mantle reservoir was predominantly removed? (2) How did the lithospheric thinning process proceed? Because of the paucity of Mesozoic mantle xenoliths, many geologists have carried out geochemical investigations on the widespread late Mesozoic mafic igneous rocks in the NCB (e.g. Jahn et al. 1999; Fan et al. 2001, 2004; Guo et al. 2001, 2003, 2004, 2005a, b, 2007; Shao et al. 2001; Zhou et al. 2001; Zhang et al. 2002, 2003, 2004; Wilde et al. 2003; Chen et al. 2004; Li et al. 2004; Yang et al. 2004; Wang et al. 2005, 2006; Zhao et al. 2005). On the whole, all these late Mesozoic mafic rocks show arc-like trace element features with large ion lithophile element (LILE) and light REE (LREE) enrichments and high field strength element (HFSE) depletion, and low radiogenic Nd with variably radiogenic Sr isotopic compositions. Such features are widely considered to be inherited from LILE- and LREE-enriched mantle sources (e.g. Jahn et al. 1999; Fan et al. 2001, 2004; Guo et al. 2001, 2003, 2004, 2005a; Zhang et al. 2002, 2003, 2004; Li et al. 2004; Wang et al. 2005, 2006). Nevertheless, there still exist chemical and isotopic differences between rocks from the various tectonic units. In this paper, we compile the available data for major and trace elements and Sr–Nd isotopes of these mafic rocks from the various tectonic units of the NCB. Our aim is to: (1) understand the chemical and isotopic heterogeneity of the SCLM and the possible enrichment processes; and (2) propose a tectonic model to interpret the extensional tectonics that were responsible for the extensive magmatism in the NCB.

Geological background The eastern NCB is bordered by the Triassic Qinling–Dabie –Sulu collisional orogen in the south and by the late Palaeozoic to early Mesozoic Central Asian Orogenic Belt (CAOB) in the north (Fig. 1; e.g. Davis et al. 2001). To the east is the NE Asian active continental margin related to the Pacific Ocean subduction. It is separated from the western NCB by the central high-pressure granulite zone, which formed in the early Proterozoic (Zhao 2001). The NCB has experienced a multiplestage crustal evolution history; the Archaean to early Proterozoic was the most important (e.g. Liu et al. 1992; Zhai et al. 2000). High-grade metamorphic rocks are widely exposed, emplaced by

extensive hypersthene granites and tonalite– trondhjemite –granites of 2.4–2.5 Ga (e.g. Zhai et al. 2000). At c. 1.8 Ga, collision between the western and eastern blocks formed a north–southtrending high-pressure granulite belt (e.g. Zhao 2001). The Archaean to early Proterozoic tectonic evolution created the framework of the metamorphic basement. A Mesoproterozoic passive rift (the Yan-Liao Rift) in the northern part of the NCB is characterized by carbonate and clastic terrestrial sediments tens of kilometres thick (e.g. the Changcheng, Jixian and Qingbaikou Systems; Hebei Bureau of Geology and Mineral Resources 1988). The rift had closed by the end of the Mesoproterozoic. From the Neoproterozoic to early Palaeozoic the NCB was a passive continental margin, where neritic carbonates thousands of metres thick were deposited. The eruption of diamond-bearing kimberlites and their mantle xenoliths suggests that the NCB was a tectonically stable block in the early Palaeozoic. A gap in sedimentation from the Late Ordovician to the Early Carboniferous marks a major period of uplift of the continent. In the Permo-Triassic, collision between the NCB and the Mongolian Block created the southern segment of the CAOB (Se¨ngor et al. 1993; Xiao et al. 2003), and the Triassic collision between the NCB and Yangtze Block formed the Qinling –Dabie –Sulu orogen in the south (e.g. Li et al. 1993). The early Mesozoic geology of the NCB is characterized by intracontinental compression and orogeny (e.g. Li 1994; Davis et al. 2001). Several workers have suggested that in the northern and southern NCB, sporadic emplacement of mafic to felsic magmas along the lithosphere-scale faults occurred as a consequence of local lithospheric extension or slab breakoff (Davis et al. 2001; Chen et al. 2003; Li et al. 2004; Guo et al. 2007). In contrast, magmatic activity was absent in the NCB interior. In the late Mesozoic lithospheric extension was dominant across the NCB, accompanied by emplacement of voluminous magmas of ultramafic –mafic to felsic composition, extensive gold mineralization, and formation of NE–SW- to NNE– SSW-trending fault-bound sedimentary basins. Such a sequence of geological events has been widely attributed to thinning of the NCB lithosphere either by removal of a lithospheric root beneath the main orogens or by convective thinning of ancient subcratonic lithospheric mantle (e.g. Fan et al. 2001, 2004; Guo et al. 2001, 2003; Xu 2001; Zhang et al. 2004). The lithospheric thinning reached its maximum in the early Tertiary, especially around the Bohai Bay, where asthenosphere-derived tholeiitic basalts were erupted; Xu et al. (1995) estimated that the

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Fig. 1. A simplified geological map of the North China Block, showing the distribution of the late Mesozoic mafic rocks. The isotopic ages of the late Mesozoic mafic rocks are from Table 1.

lithospheric thickness was 55 –70 km. Thermal decay of upwelling asthenosphere and lithospheric accretion took place as a result of melt extraction, and led to the Neogene –Quaternary extrusion of widespread xenolithic alkaline basalts.

Geochronology of Mesozoic mafic magmatism in NCB We subdivide the NCB into three tectonic units from north to south: (1) the northern unit, which includes the Yinshan–Yanshan –Liaoxi tectonomagmatic belt, where Mesozoic magmatic activity began in the early Jurassic (e.g. Davis et al. 2001; Li et al. 2004; Guo et al. 2007); (2) the interior,

mainly in the Taihangshan and Luxi areas; (3) the southern margin, which includes regions along the Triassic Qingling– Dabie– Sulu collisional orogen (Fig. 1b). Widespread late Mesozoic mafic rocks in the NCB include (Fig. 1) intrusive ultramafic – mafic complexes, (e.g. Jinan, Zouping and Taihanshan in the NCB interior; Guo et al. 2001; Zhang et al. 2004; Wang et al. 2005b), the Dabie terrane in the southern NCB (Jahn et al. 1999; Li et al. 1999; Zhao et al. 2005), and mafic dykes and lavas all over the NCB (Fan et al. 2001, 2004; Shao et al. 2001; Zhou et al. 2001; Guo et al. 2004, 2005a, b; Yang et al. 2004; Wang et al. 2005). Available K –Ar, Ar –Ar and U –Pb zircon dating of these mafic rocks indicates that their emplacement age

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spanned a range from 140 to 105 Ma with a peak around 125 Ma. The isotopic ages of late Mesozoic mafic rocks in the NCB are summarized in Table 1.

Petrology of late Mesozoic mafic rocks in the NCB The late Mesozoic NCB mafic rocks comprise a wide spectrum of rock types, including ultramafic cumulates (e.g. pyroxenites), gabbros, diorites, lamprophyres, basalts, trachybasalts, basaltic andesites and trachyandesites. Their country rocks comprise the Archaean–early Proterozoic metamorphic basement, Triassic HP –UHP metamorphic terranes, and Mesoproterozoic to Mesozoic sediments. A cumulate texture is characteristic of the ultramafic–mafic intrusions, such as in the Jinan norites and the pyroxenites in the Dabie terrane. Most mafic extrusive rocks, such as the mafic lavas and dykes and lamprophyres, show a porphyritic texture with predominant phenocrysts of clinopyroxene, plagioclase and subordinate olivine, and amphibole in basaltic andesites and trachyandesites. The gabbros and gabbroic diorite generally show a mosaic texture with euhedral to subhedral shapes of the major phases such as plagioclase, clinopyroxene and amphibole. The main petrographic features of the late Mesozoic NCB mafic rocks are summarized in Table 2.

Geochemistry of late Mesozoic mafic rocks in the NCB Major elements On a loss on ignition-free basis, the late Mesozoic NCB mafic rocks have large ranges of SiO2 (41– 57%) and MgO (2.5–19%), with lower MgO in the northern NCB rocks. Some ultramafic samples from the southern region and interior of the NCB that have a cumulate texture (e.g. norites from the Jinan intrusion and pyroxenites from the Dabie terrane (Table 2; Jahn et al. 1999; Guo et al. 2001; Zhao et al. 2005) are mafic cumulates. These rocks generally have higher MgO and CaO, and lower Al2O3, K2O, Na2O and P2O5 (Fig. 2). Strictly, cumulates cannot represent mantle-derived magmas but are assemblages of early fractional phases, and their elemental features are mainly controlled by the major minerals. Therefore, we will avoid using these rocks to discuss the magmatic processes and source characteristics in the following text. All late Mesozoic NCB mafic rocks show a general increase in SiO2, Al2O3, K2O and P2O5, and decrease in FeOT, CaO and CaO/Al2O3

following a decrease in MgO. Rocks from the southern region and interior of the NCB also show an increase in Na2O and TiO2, whereas the northern NCB samples show random variation in Na2O and TiO2. At a given MgO the northern NCB rocks generally have higher TiO2 and P2O5 than the others, and the NCB interior rocks have the lowest TiO2 and P2O5.

Trace elements To avoid effects caused by fractional crystallization and/or crustal contamination, we selected those samples free of a cumulate texture to calculate the average abundances of mafic rocks from each locality. Primitive mantle-normalized trace element spidergrams of mafic rocks from various localities are shown in Figure 3. All mafic rocks from the NCB have LILE and LREE enrichments but HFSE depletion, similar to modern arc basalts or contaminated continental flood basalts (e.g. Grove & Kinzler 1986; Peng et al. 1994), but completely different from asthenosphere-derived magmas (e.g. MORB and OIB; Fig. 3; Sun & McDonough 1989). There are still some trace element differences between the mafic rocks from the various units of the NCB. For instance, the NCB interior rocks have lower La, at about 40 times that of primitive mantle, whereas other samples from the northern and southern NCB span a range of La that is 80 – 200 times that of primitive mantle. Additionally, the NCB interior samples show a trough in Zr–Hf with respect to Sm and Eu, whereas samples from the northern NCB exhibit Th –U depletion relative to Ba and La.

Sr – Nd isotopes Sr– Nd isotope data of the late Mesozoic NCB mafic rocks are listed in Table 3 and plotted in Figure 4. All rocks have variably radiogenic Sr but unradiogenic Nd isotopic compositions, spanning a range of 87Sr/86Sr(i) from 0.7040 to 0.7104 and 1Nd(t) from 219.2 to 25.4, completely different from that of the Cenozoic basalts (87Sr/86Sr(i) ¼ 0.703 2 0.705 and 1Nd(t) ¼ 23 to þ6) in eastern China (e.g. Liu et al. 1994). In addition, there are considerable Sr– Nd isotopic differences among rocks from the various tectonic units. In general, the southern NCB rocks have the highest 87Sr/86Sr(i) and lowest 1Nd(t) values, whereas the NCB interior samples have the lowest 87Sr/86Sr(i), but relatively lower 1Nd(t) than the northern NCB rocks. Nevertheless, even samples from the same tectonic unit also show fairly variable Sr–Nd isotopic compositions (Table 2 and Fig. 4).

Southern NCB Fangcheng North Huaiyang Dabie terrane Dabie terrane Laiyang basin Jiaodong

Zouping Zouping Jiyang basin Laiwu Taihangshan Northern NCB Liaoxi Yanshan Yanshan Hannuoba

NCB interior Jinan

Locality

Olivine basalt Mafic volcanic rocks Ultramafic–mafic intrusions Mafic dykes Mafic volcanic rocks Lamprophyres and mafic dykes

Mafic volcanic lavas Mafic dykes Mafic volcanic lavas Mafic granulites and metagabbros

Hyperite – norite–gabbro– diorite intrusion Gabbro – diorite intrusion Basaltic trachyandesite Basalts and andesites Gabbro and diorite Gabbro and diorite

Rock type

125 132 – 116 122 – 125 132 – 125 105 – 112 121 – 125

135 – 120 124 136 – 115 140 – 120

115 128 – 130 136 – 107 120 122 – 125

115

Age (Ma)

Table 1. A summary of isotopic ages of the NCB late Mesozoic mafic rocks

K– Ar K– Ar Zircon, U – Pb, Ar – Ar Ar– Ar K– Ar K– Ar

K– Ar, Ar– Ar K– Ar K– Ar, Ar– Ar Zircon, U – Pb

Ar– Ar K– Ar K– Ar Rb – Sr Zircon, U – Pb

Ar– Ar

Method

Zhang et al. 2002 Fan et al. 2004 Jahn et al. 1999; Li et al. 1999; Zhao et al. 2005 Wang et al. 2005a Guo et al. 2005a Guo et al. 2004; Yang et al. 2004

Chen et al. 1997 Shao et al. 2001 Davis et al. 2001 Fan et al. 1998; Wilde et al. 2003; Liu et al. 2004

Lin et al. 1996 Guo et al. 2003 This study Chi et al. 1994 Wang et al. 2005b

Lin et al. 1996

References LATE MESOZOIC MAFIC MAGMATISM 81

Mesozoic sediments Hosted by Cenozoic basalts

Ol þ Cpx þ Pl Opx þ Cpx þ Pl þ Q

Cpx þ Pl þ amph Ol þ Cpx þ Pl þ amph Ol þ Cpx þ Pl þ Kf þ Bi þ Amph

Porphyritic fabrics Porphyritic fabrics Porphyritic fabrics

Basaltic trachyandesite Trachybasalt and trachyandesite Lamprophyre

Ol þ Opx þ Cpx þ Pl þ Amph þ accessories Cpx þ Pl þ amph

Mesozoic sediments

Cumulate texture and gabbroic fabrics Porphyritic fabrics

Porphyritic fabrics Granulitic and cumulate texture

Mesozoic sediments Mesozoic sediments

Cpx þ Plþ Amph Ol þ Cpx þ Pl þ Bi Cpx þ Pl þ amph

Pyroxenite, gabbro and diorite Mafic dyke

Trachybasalt and basaltic trachyandesite High-Mg andesite Mafic granulite and metagabbro

Mesozoic sediments

Ol þ Cpx þ Pl

Metamorphic terranes

North Dabie complex, Mesozoic sediments Mesozoic sediments Mesozoic sediments

Dabie metamorphic complex

Palaeozoic carbonates Palaeozoic carbonates Metamorphic basement

Palaeozoic carbonates

Basement rocks

Ol þ Opx þ Cpx þ Pl þBi þ accessories Ol þ Opx þ Cpx þ Pl þ Bi þ Q þ accessories Ol (minor) þ Cpx þ Pl þ Amph Ol þ Cpx þ Pl þ amph þ accessories

Mineral assemblage/phenocryst

Ol, olivine; Opx, orthopyroxene; Cpx, clinopyroxene; Amph, amphibole; Bi, biotite; Pl, plagioclase; Kf, potassium feldspar; Q, quartz. Data sources are the same as in Table 1.

Sulu belt

Southern NCB Dabie terrane

Liaoxi Hannuoba

Jibei

Porphyritic fabrics

Porphyritic fabrics Porphyritic fabrics

Basaltic trachyandesite Mafic dyke

Basalt and andesite

Basaltic trachyandesites Gabbro and diorite

Taihangshan

Jiyang Basin Northern NCB Xishan

Porphyritic texture Mosaic texture and gabbroic fabrics Porphyritic fabrics

Gabbro and diorite

Zouping

Texture Gabbroic fabrics and cumulate texture Gabbroic fabrics

Rock type

Norite – gabbro– diorite

NCB interior Jinan

Locality

Table 2. A summary of petrographic features of the late Mesozoic NCB mafic rocks

82 W.-M. FAN ET AL.

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Fig. 2. MgO v. major element oxide (in wt%) (a– h) and CaO/Al2O3 (i) and Ni (in ppm) (j) diagrams of the late Mesozoic NCB mafic rocks. Data for the NCB interior are from Guo et al. (2001, 2003), Chen et al. (2004), Zhang et al. (2004) and Wang et al. (2005b); for the northern NCB from Shao et al. (2001), Zhou et al. (2001) and Guo et al. (2007); and for the southern NCB from Jahn et al. (1999), Fan et al. (2001, 2004), Zhang et al. (2002), Guo et al. (2004, 2005a), Yang et al. (2004), Wang et al. (2005a) and Zhao et al. (2005).

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Fig. 3. Primitive mantle-normalized trace element spidergrams of the late Mesozoic mafic rocks from different tectonic units of the NCB. Normalized values for primitive mantle are from Sun & McDonough (1989). Data sources are same as in Figure 2. INCB, NNCB and SNCB refer to the interior, northern and southern NCB.

18

8

30

25

10 6

Jiyang basin

Southern NCB North Dabie terrane

Sulu belt

Jiaodong Fangcheng

9

Yanshan

20

18

Hannuoba

NCB interior Jinan – Zouping Taihangshan

18

Sample number

Northern NCB Liaoxi– Jibei

Tectonic unit

Mafic intrusions, dykes and basaltic lavas Basaltic lavas and lamprophyres Mafic dykes Olivine basalts

Mafic intrusions and mafic lavas Gabbros and gabbroic diorites Basalts and basaltic andesites

Mafic lavas and high-Mg andesites Mafic granulites and metagabbros Mafic lavas and dykes

Lithology

Sr/86Sr(i)

87

0.7075– 0.7114 0.7096– 0.7101

0.7072– 0.7091

0.7070– 0.7092

0.7041– 0.7049

0.7049– 0.7054

0.7042– 0.7058

0.7053– 0.7069

0.7059– 0.7086

0.7061– 0.7069

Table 3. Sr – Nd isotope compositions of the NCB late Mesozoic mafic rocks

Zhang et al. 1998; Zhou et al. 2002 Shao et al. 2001; Zhou et al. 2001; Guo et al. 2007

218.2 to 29.0 214.1 to 27.2

Fan et al. 2001; Guo et al. 2004, 2005a Yang et al. 2004 Zhang et al. 2002

218.2 to 215.5 217.2 to 211.0 213.1 to 214.2

213.9 to 25.4

Jahn et al. 1999; Fan et al. 2004; Wang et al. 2005b

Our unpublished data

216.6 to 211.2

219.2 to 213.9

Guo et al. 2001, 2003; Zhang et al. 2004 Chen et al. 2004; Wang et al. 2005a

218.7 to 26.1

Zhou et al. 2001; Zhang et al. 2003; Guo et al. 2007

References

29.0 to 26.0

1Nd(t)

LATE MESOZOIC MAFIC MAGMATISM 85

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Fig. 4. Sr– Nd isotope data of late Mesozoic mafic rocks from the NCB. Data for the Cenozoic NCB sediments are our unpublished data; data for granulite-facies xenoliths are from Zhang et al. (1998), Zhou et al. (2002), Huang et al. (2004) and Li et al. (2004); and those for the Cenozoic basalts of eastern China are compiled from Liu et al. (1994). The other data sources are same as in Figure 2. UCC, upper continental crust; LCC, lower continental crust.

Chemical and isotopic heterogeneity of the SCLM beneath the NCB Possible effect of crustal contamination Before using the available geochemical data to constrain the source characteristics of the melted mantle sources, it is necessary to evaluate the effect of crustal contamination or assimilation – fractional crystallization (AFC) during magma evolution, because all late Mesozoic NCB mafic rocks have significant negative Nb –Ta anomalies in primitive mantle-normalized spidergrams, radiogenic Sr and unradiogenic Nd isotopic compositions. Mass balance considerations suggest that such features cannot be realistically attributed to crustal contamination or AFC of asthenospherederived magmas such as MORB and OIB. As shown by Figure 5a, addition of 40–60% upper continental crust (UCC) or lower continental crust (LCC) to OIB-type magma would produce the observed Sr –Nd isotopic compositions in the NCB mafic rocks. Introduction of such high proportion of crustal material could not be re-equilibrated with a heat budget. The Sr– Nd isotopic compositions in these rocks could not be produced by assimilation by UCC or LCC (Fig. 5a). For instance, with a remaining magma fraction less than 0.1, LCC assimilation will create a magma with 1Nd(t) around 25. Such a low remaining magma fraction

would lead to extensive magma fractionation and the created magma should have been felsic rather than mafic. Although some continental basalts with arc-type trace element features (e.g. some contaminated continental flood basalts and continental potassic mafic magmas) could be produced by interaction between asthenosphere-derived magma and the overlying enriched lithospheric mantle (e.g. Arndt & Jenner 1986; Peng et al. 1994; O’Brien et al. 1995; Shimizu et al. 2005), the absence of contemporaneous MORB- or OIB-type mafic magmas across the block tends to rule out this possibility. We thus consider that the geochemical features in the NCB mafic rocks were mainly derived from enriched lithospheric mantle. The Sr isotope ratios in the southern and northern NCB mafic rocks are generally higher than those in rocks from the NCB interior: such a difference could have been produced by crustal contamination or AFC. We selected the possible endmember components and partition coefficients (KD) for Sr and Nd to model this possibility. As shown by Figure 5b, for a magma with 87Sr/ 86 Sr(i) at 0.704 and 1Nd(t) at 27, an addition of .50% UCC into such a magma would produce a new melt similar to the lowest 87Sr/86Sr(i) (c. 0.707) in the southern NCB rocks. AFC modelling results show that it is necessary for the remaining magma fraction of ,0.2 to produce a magma with 87 Sr/86Sr(i) c. 0.707. Neither LCC contamination

LATE MESOZOIC MAFIC MAGMATISM

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Fig. 5. Sr– Nd isotopic modelling results of the effect of crustal contamination and AFC on the late Mesozoic NCB mafic rocks. In (a) is shown crustal contamination and AFC of an assumed OIB-type basalt, and of a hypothetical Nd primary basalt from the NCB interior. The assimilation rate is assumed to be 0.25, KSr D ¼ 1.2, KD ¼ 0.5. Dashed lines represent AFC curves and continuous lines denote bulk mixing curves between the hypothetical primary basalt and crustal components. The NCB UCC composition is compiled from the Cenozoic NCB sediments (our unpublished data), and the NCB LCC composition is compiled from data for the granulite xenoliths and Archaean high-grade metamorphic terranes in the NCB of Jahn & Zhang (1984), Zhang et al. (1998), Zhou et al. (2002), Huang et al. (2004) and Liu et al. (2004). The hypothetical primary basalt is a basaltic sample from the Jiyang basin in the NCB interior (our unpublished data). The composition of OIB is arbitrarily assumed. Tick marks on the AFC curves represent increments of 0.1 of the remaining magma fraction, and those on the bulk mixing curves denote the percent of the mixed crustal component. The other calculation parameters used for Sr–Nd isotopic modelling are listed below:

88

Component OIB UCC LCC Hypothetical primary basalt

W.-M. FAN ET AL.

Sr (ppm)

Nd (ppm)

800 210 350 923

35 40 20 27

nor assimilation can produce the Sr isotopic difference between the interior and the southern NCB rocks. The northern NCB mafic rocks, which have slightly higher 87Sr/86Sr(i) than the NCB interior samples, could not be simply interpreted as the contamination or AFC products of the NCB interior mafic magmas. Sr–Nd isotopic modelling results suggest that either an addition of 30% UCC into a magma with 87Sr/86Sr(i) at 0.704 or UCC assimilation with the remaining magma fraction of around 0.4 could create the lowest 87 Sr/86Sr(i) (c. 0.7055) sample from the northern NCB (Fig. 3b). Nevertheless, the northern NCB mafic rocks show considerable Th –U depletion relative to Ba and La; some samples even have a Th/La ratio lower than that of N-MORB (Guo et al. 2007). Such trace element features call for either contamination or assimilation by lower crustal rocks (Rudnick & Fountain 1995). Increase of 87Sr/86Sr(i) from 0.704 to 0.7055 required either addition of .50% LCC to the hypothetical primary magma or LCC assimilation with the remaining magma fraction of ,0.1. Similarly, both mass balance and heat considerations argue against an origin of these melts by LCC contamination or assimilation of the primary melts from the NCB interior. Accordingly, the difference in Sr –Nd isotopes between the mafic rocks from different units was mainly inherited from heterogeneous mantle sources. The effect of crustal contamination and AFC on the mafic rocks from individual localities has been discussed in previous studies (e.g. Jahn et al. 1999; Fan et al. 2001, 2004; Guo et al. 2001, 2004, 2005a, b, 2007; Zhou et al. 2001; Zhang et al. 2002, 2003, 2004; Yang et al. 2004; Wang et al. 2005, 2006; Zhao et al. 2005). Using various methods, they proved that the majority of the NCB mafic rocks had escaped significant crustal contamination through their extrusion or intrusion. In the following discussion, we will select those samples free of contamination and cumulate texture and having MgO .5% from the southern region and interior of the NCB. For the purpose of comparison and because of the existence of only a few samples with MgO .5% in the northern NCB, the rocks with a MgO range of 4–5% from this unit are also used for discussion.

87

Sr/86Sr(i) 0.703 0.716 0.708 0.704

1Nd(t) þ8 218 230 27

Major element constraints The majority of samples from both the southern region and interior of the NCB show an increase of SiO2, Al2O3, K2O, Na2O, TiO2 and P2O5, and a decrease of FeOT, CaO, CaO/Al2O3 and Ni following magma differentiation, indicating a dominant fractionation of olivine and clinopyroxene (Fig. 2). However, the northern NCB rocks show an increase of SiO2, Al2O3, K2O and P2O5, and a decrease of TiO2, FeOT, CaO, CaO/Al2O3 and Ni following magma evolution, suggesting a fractional assemblage of olivine þ clinopyroxene + amphibole + ilmenite. An approach commonly used for comparison of large mafic rock datasets is to extrapolate the less evolved rocks back to 8% MgO using a best-fit linear regression. Using this method, Klein & Langmuir (1987) first identified the global trends of MORB to estimate the melting conditions, such as pressure, temperature and degree of partial melting. This has been widely applied to continental flood basalt petrogenesis and yielded important information about melting conditions and source characteristics (e.g. Turner & Hawkesworth 1995; Garland et al. 1996; Xu et al. 2001). Extrapolation of the late Mesozoic NCB mafic rocks to 8% MgO was carried out by fitting a simple correlation line, following the method and rationale of Turner & Hawkesworth (1995). Only samples with a MgO range of 5–10% from the interior and southern NCB and those with MgO .4% from the northern NCB are regressed. The correlations between MgO and SiO2, FeOT, TiO2 and Na2O of the late Mesozoic mafic rocks from different units of the NCB are illustrated in Figure 6, and the calculation results of Si8, Fe8, Na8 and Ti8 are shown in Figure 7. The northern NCB samples have the highest Fe8, Ti8 and Na8 and the lowest Si8 of the NCB mafic rocks, and they show negative correlations between Si8, Fe8 and Ti8. In contrast, the NCB interior samples have the lowest Ti8 and Na8 and the highest Si8, and the southern NCB mafic rocks have Si8, Fe8, Ti8 and Na8 between those of the NCB northern and interior rocks. Experimental melting results show that the Fe content of a melt is positively correlated with

Fig. 6. Correlations between MgO and SiO2, FeOT, Na2O and TiO2 of the late Mesozoic mafic rocks from different units of the NCB. (a –d) The northern NCB rocks; (e–h) the southern NCB rocks; (i– l) the NCB interior rocks. Data sources are as in Figure 2.

LATE MESOZOIC MAFIC MAGMATISM 89

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pressures. Comparing the experimental results (e.g. Kushiro 1990; Hirose & Kushiro 1993; Baker & Stolper 1994), the highest Fe8 and lowest Si8 in the northern NCB mafic rocks suggest their derivation from high-pressure melting, and the highest Ti8 and Na8 could be caused by low degrees of partial melting. However, a different melting condition is not a unique interpretation for the compositional distinction between the mafic rocks from different units. Melt generation from a refractory mantle source can lower Fe8, and the Ti8 gap between the northern NCB rocks and the other samples from the NCB suggests that the northern NCB mafic magmas were derived from a more fertile mantle source than that from which the southern and interior NCB melts were derived (Fan & Guo 2005). The distinction between these alternatives is not straightforward and will be addressed in the following sections.

Trace element constraints

Fig. 7. Si8 v. Fe8, Ti8 and Na8 plots of the late Mesozoic NCB mafic rocks. Numbers of samples: 14 from the northern NCB; 31 from the NCB interior; 39 from the southern NCB.

pressure and source composition, and thus melts either at high pressures or derived from a fertile source will have high Fe (e.g. Kushiro 1990; Hirose & Kushiro 1993; Baker & Stolper 1994). In contrast to Fe, the Si content is negatively correlated with melting pressure, and it increases in a water-saturated system. The Ti content of a melt is a means of gauging source fertility, as it is incompatible during melting and sensitive to melt removal. Na mainly reflects the degree of melting as it acts as a moderately incompatible element in mantle conditions. However, Na is less incompatible, as it is retained in the residue and incorporated in the jadeite component of clinopyroxene at high

As previously described, some late Mesozoic mafic rocks from different tectonic units have trace element differences. Here we present a comparison to uncover possible differences between the mantle sources from which these magmas were derived. Because all late Mesozoic NCB mafic rocks discussed here had a predominant fractional assemblage of olivine and clinopyroxene, trace elemental ratios such as LILE/LREE, Th/U, Th/La and Th/ Nb changed only slightly, because of the low and similar partition coefficients of these elemental pairs between the fractional phases and basaltic melts; nor would partial melting change these elemental ratios significantly. The key trace element features of the late Mesozoic mafic magmas are listed in Table 4 and plotted in Figure 8. As shown by Figure 8, the northern NCB rocks are characterized by the lowest Th/La, Th/Nb and La/Nb, and moderate La and La/Sm. The NCB interior rocks have higher Ba/La and Sr/Nd but lower La, La/Sm, (Zr/Sm)N and (Hf/Sm)N than the others in the NCB, whereas the southern NCB samples have the highest Th, U, Th/U, La and La/Sm, and the largest range of La/Nb and Th/Nb. We thus consider that such features were mainly inherited from the melting sources. Irrespective of the effect caused by degrees of melting, one may speculate that the melting source for the NCB interior rocks also had low LREE contents. The lowest Th/La and Th/Nb ratios in the northern NCB rocks suggest that they were derived from a low-Th source, whereas the southern NCB samples, which have the highest Th, U, La and La/Sm, probably originated from a source highly enriched in these elements.

0.24 –1.53 (0.71)

2.0 – 14 (5.5)

*Values in parentheses are averages of the trace element ratios.

0.22 –1.1 (0.54) 0.08 –0.16 (0.14)

1.8 – 7.6 (4.2) 1.5 – 3.8 (2.7)

NCB interior (n ¼ 39) Northern NCB (n ¼ 13) Southern NCB (n ¼ 53)

Th/Nb

La/Nb*

Tectonic unit

3.1– 7.7 (5.6)

2.8– 5.9 (3.9) 2.7– 6.4 (4.6)

Th/U

0.08– 0.33 (0.13)

0.07– 0.25 (0.13) 0.04– 0.08 (0.05)

Th/La

Table 4. Key trace element features of the late Mesozoic NCB mafic magmas

8.3– 70 (31)

20 – 264 (65) 13 – 65 (28)

Ba/La

7.8– 27 (17)

18 – 64 (34) 8.0– 27 (17)

Sr/Nd

4.9– 9.1 (7.2)

2.7– 5.3 (4.3) 4.8– 8.0 (6.0)

La/Sm

0.47– 1.51 (0.97)

0.20– 1.00 (0.71) 0.81– 1.28 (1.03)

(Zr/Sm)N

0.25– 1.01 (0.69)

0.41– 1.16 (0.83) 0.70– 1.07 (0.89)

(Hf/Sm)N

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Fig. 8. Trace element plots of the NCB late Mesozoic mafic rocks. In (a), the southern NCB rocks have the highest Th and U, whereas samples from the interior and northern NCB have comparable Th and U contents. In (b) and (c), the NCB interior samples have the highest Ba/La and Sr/Nd ratios, and the northern NCB rocks have the lowest Th/La. In (d), the northern NCB rocks have lower Th/Nb and La/Nb compared with samples from the interior and southern NCB. In (e), the NCB interior rocks have the lowest La concentrations of the three groups and the highest La and La/Sm in the southern NCB samples. In (f), the NCB interior rocks have subchondritic (Zr/Sm)N and (Hf/Sm)N ratios, suggesting an important role of carbonatite metasomatism in the melting sources. Data sources are the same as in Figure 2.

Sr –Nd isotopic variations across the NCB To show possible Sr–Nd isotopic variations across the NCB, we here compile 142 sets of Sr–Nd

isotope data for the late Mesozoic mafic rocks in the NCB from the literature and our unpublished data from the Jiyang basin around the Bohai Bay. The detailed profiles are illustrated in Figure 9.

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Fig. 9. Spatial Sr and Nd isotopic variations across the NCB, showing highly heterogeneous Sr and Nd isotopic compositions between different tectonic units and even between rocks from the same locality.

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There is a clear Sr–Nd isotopic heterogeneity in the mantle sources beneath the different units of the NCB. For instance, the initial 87Sr/86Sr(i) ratio ranges from c. 0.7065 in Jibei –Liaoxi to c. 0.7055 in Xishan, and to c. 0.7050 in the NCB interior (e.g. Jiyang basin, Taihangshan, Jinan and Zouping). From the NCB interior to the southern NCB, 87Sr/86Sr(i) increases to a maximum (.0.709) at Fangcheng, and slightly decreases from Sulu to North Huaiyang and then to the North Dabie terrane. Similarly, 1Nd(t) shows variations across the NCB. It is highest in the Jibei – Liaoxi region, with average values of around 28.0, and is c. 213 in Xishan, c. 214 at Taihangshan, 210 at Jiyang basin and c. 215 in Jinan– Zouping, and has the lowest values of around 216.5 in the Dabie –Sulu region. Despite having the highest 87Sr/86Sr(i), the Fangcheng olivine basalts have 1Nd(t) values of c. 214, somewhat higher than for other samples from the southern NCB. The highly variable Sr–Nd isotopic compositions in the late Mesozoic mafic rocks indicate highly heterogeneous mantle sources across the NCB.

Possible enrichment processes The arc-type trace element features and enriched Sr–Nd isotopic signatures of all late Mesozoic mafic rocks from the NCB suggest that their melted mantle sources experienced LILE and LREE enrichments prior to the melting event (e.g. Menzies et al. 1987). Different types of melt or fluid metasomatism will create distinctive trace element and isotope compositions in metasomatized mantle (e.g. Menzies et al. 1987; You et al. 1996). For instance, a carbonatite melt or fluid has a preferential capacity to host LREE, Ba and Sr, a relatively low capacity for Rb and K, and the weakest capacity for HFSE and Ti. Therefore, a source having undergone carbonatite metasomatism would have low Rb/Sr and Sm/Nd, and HFSE and Ti depletion (Dupuy et al. 1992; Rudnick et al. 1993; Furman 1995). Slab-derived fluids preferentially host Rb, K, Ba, Sr and U, but have a low capacity for REE, Th and HFSE, so a mantle metasomatized by slab-derived fluids will have high U and Rb/Sr, and moderate Sm/Nd, and low Th and HFSE (e.g. You et al. 1996; Hawkesworth et al. 1997; Macdonald et al. 2000). A slab melt is characterized by TTG- or adakite-type elemental features (e.g. high Al, Na and Sr, and low Y and HREE; (Defant & Drummond 1990; Smithies 2000; Martin et al. 2005). The moderate radiogenic Sr and low radiogenic Nd isotopic compositions in the NCB interior mafic rocks are similar to those derived from the SCLM beneath Archaean cratons. Such EM1-like isotopic

signatures were probably caused by pervasive metasomatism of a carbonatitic melt or fluid with moderate Rb/Sr and low Sm/Nd (e.g. Menzies et al. 1987; Dupuy et al. 1992; Rudnick et al. 1993; Bedini et al. 1997; Guo et al. 2001; Zheng et al. 2005), as indicated by their higher Ba/La and Sr/ Nd ratios (Fig. 8b and c) in the NCB interior mafic rocks and the occurrence of contemporaneous carbonatite magmatism in the NCB interior (e.g. Ying et al. 2004). It is generally considered that carbonatitic melts or fluids have a much stronger capacity to host LREE than HFSE (e.g. Dupuy et al. 1992; Rudnick et al. 1993; Guo et al. 2001; Zheng et al. 2005); a carbonatite metasomatism can also account for the considerable Zr–Hf depletion or subchondritic (Zr/Sm)PM and (Hf/ Sm)PM ratios in these rocks (Figs 3a and 8f). The northern NCB mafic rocks show spatial Sr – Nd isotopic variations. At Xishan, near the NCB interior, the mafic rocks have similar Nd isotopic ratios, and higher 87Sr/86Sr(i) ratios than those from the NCB interior, implying a recent LILE enrichment event prior to magma generation. In the Jibei –Liaoxi area, farther from the NCB interior, mafic rocks have higher 87Sr/86Sr(i) and 1Nd(t), signatures of subduction-related enrichment processes (e.g. You et al. 1996; Macdonald et al. 2000). Considering its high Fe8, Ti8 and Na8, low Si8, and remarkable Th –U depletion, we consider that the lithospheric mantle had been modified by recycled ancient lower crustal materials with low Th/La, probably during the late Palaeozoic –early Mesozoic Palaeo-Asian Ocean subduction (Li et al. 2004; Guo et al. 2007). The southern NCB mafic rocks have the highest 87 Sr/86Sr(i) and lowest 1Nd(t) of the rocks across the NCB. Such isotopic signatures suggest that the lithospheric mantle had been modified by melts or fluids with high Rb/Sr and low Sm/Nd ratios related to the Triassic collision between the NCB and Yangtze Block (Amelin et al. 1996; Jahn et al. 1999; Fan et al. 2001; Zhang et al. 2002; Guo et al. 2004, 2005b; Wang et al. 2005). To account for the highest Th, U and LREE concentrations in the southern NCB mafic rocks, we prefer a process of enrichment by melt metasomatism rather than fluid metasomatism, because slabderived fluids have a low capacity to transfer Th. Based on superchondritic Zr/Nb and Th/U ratios in the mafic rocks from the Sulu belt, Guo et al. (2005a) pointed out that the melts were probably derived from diorite–trondhjemite –tonalite–granitic gneisses like those of the Kongling Group of the subducted Yangtze lithosphere (Gao et al. 1999; Ma et al. 2000). On the other hand, the moderate Nb/Ta and Zr/Hf fractionations in the majority of the southern NCB mafic rocks suggest that this enrichment process occurred at high pressures with a

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rutile- and CO2-rich metasomatic assemblage (e.g. Guo et al. 2004, 2005b). In summary, distinctive and diagnostic enrichment processes were responsible for the chemical and isotopic heterogeneity beneath the different tectonic units of the NCB.

Implications for the lithospheric thinning of the NCB The late Mesozoic emplacement of voluminous mafic melts across the NCB is unique in global cratonic regions. Both numerical modelling and thermal considerations suggest that such extensive melting of lithospheric mantle requires either a voluminous heat supply or significant lithospheric extension, or both. However, evidence from geophysical surveys, the arc-like trace element patterns, and the enriched Sr–Nd isotopic signatures in these rocks argue against the presence of a plume beneath the NCB. Therefore, we attribute the extensive late Mesozoic magmatism to decompressional melting of enriched mantle sources as a consequence of lithospheric extension and thinning. The causes of the early Cretaceous lithospheric extension of the NCB are still controversial. Wilde et al. (2003) correlated the lithospheric thinning event with the breakup and dispersal of Gondwanaland by a major mantle overturn, fuelled by the destruction of oceanic lithosphere and triggered by its sinking into the lower mantle during the subsequent accretion of Asia. In a lateral extrusion model, Guo et al. (2003) considered that tectonic forces from the compression of the Siberian plate in the north and from the Tethyan belt in the SW, together with pull-apart stresses from the rapid northward movement of the palaeo-Pacific Ocean, would have triggered lithospheric extension and formation of the regional NE–SW- to NNE– SSW-trending structure. Numerical modelling results suggest that significant thinning of a lithosphere requires a high heat flux from the convective asthenosphere, whereas tectonic stretching plays only a limited role in lithospheric attenuation (e.g. Lin et al. 2005). Lin et al. pointed that the lithospheric thinning and extensive tectonomagmatic activities in the NCB reflected the thermotectonic responses to the high heat flow responsible for the Cretaceous superplume event (Larson 1991). For instance, the production of oceanic crust in the Phanerozoic reached its peak during the early Cretaceous, and this defined a long Cretaceous normal magnetic period from 125 to 90 Ma, consistent with the eruption time of the Ontong– Java large igneous province in the southwestern Pacific Ocean (e.g. Neal et al. 1997). Larson (1991) proposed a superplume hypothesis to interpret the

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global Cretaceous event, and Filatova (1998) suggested that the formation of Cretaceous active continental margins was correlated with this global event. Tectonic reconstruction of the eastern Asian continent indicates that the NCB was strongly affected by surrounding plate interactions (e.g. Engebretson et al. 1985; Faure & Natal’in 1992; Zhao et al. 1994; Davis et al. 2001). On the one hand, the NCB was dragged northward by the rapid northward movement of the palaeo-Pacific plate (e.g. Faure & Natal’in 1992; Zhao et al. 1994; Engebretson et al. 1985), which exerted a NNE –SSW-trending pull-apart stress on the NCB and its surrounding regions. Such a shearing stress led to the formation of pull-apart basins and strikeslip extension and collapse of the lithosphere (e.g. Xu et al. 1987; Ruppel 1995; Fan et al. 2001). On the other hand, the NCB was simultaneously subjected to southward and northeastward compression from the Siberian plate and the Tethyan tectonic domain, and possibly from the southern Indo-China Block, triggering a northeastward lateral escape of the NCB lithosphere (e.g. Menzies et al. 1993; Yin & Nie 1993; Guo et al. 2003; Fan & Guo 2005). The resulting tectonics, together with the global high heat flux in the early Cretaceous, caused extensive lithospheric extension and thinning of the NCB and the subsequent large-scale melting of the mantle sources (Fig. 10). As a whole, we consider that the principal late Mesozoic tectonomagmatic activities in the NCB were the intracontinental responses to surrounding plate interactions.

Fig. 10. A proposed tectonic model to interpret the extensive late Mesozoic tectonomagmatic activities in the NCB. This demonstrates the effects caused by surrounding plate interactions. (See details in the text.)

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Summary Available geochemical data from the NCB late Mesozoic mafic rocks indicate that the SCLM was chemically and isotopically heterogeneous beneath different tectonic units. Such differences resulted from distinctive enrichment processes before the magma generation. The NCB interior mafic rocks have the lowest Fe8, Na8, Ti8 and LREE, the highest Si8, HFSE depletion relative to REE, and EM1-like Sr–Nd isotopic signatures. These melts were derived from a chemically refractory mantle reservoir similar to that beneath cratonic lithospheric mantle. Pervasive carbonatite metasomatism was responsible for the LREE and LILE enrichments. In contrast, the northern NCB mafic rocks have the highest Fe8, Na8 and Ti8, lower Si8, Nb–Ta and Th –U depletion, and more radiogenic Sr and Nd isotopic compositions than those from the NCB interior. These magmas originated from a chemically fertile mantle source, which had been modified by recycled lower crustal materials, possibly related to the subduction of the palaeo-Asian Ocean. The southern NCB mafic magmas are chemically transitional between the interior and northern NCB rocks, and have the highest Th, U, LREE and 87Sr/86Sr(i), and the lowest 1Nd(t), of the NCB samples. These melts were derived from highly LILE- and LREE-enriched mantle sources modified by subducted Yangtze lower–middle crustal rocks during the Triassic collision between the NCB and Yangtze Block. The widespread late Mesozoic mafic magmatism in the NCB calls for a voluminous heat supply and significant lithospheric extension and thinning. Considering the global Cretaceous events that required a higher heat flux from the convective mantle, we consider that the early Cretaceous tectonomagmatic activities in the NCB were intracontinental responses to the surrounding plate interactions, which include pull-apart stresses induced by contemporaneous northward motion of the palaeo-Pacific plate and compression from the northern Siberian plate and the southwestern Tethyan belt, and possibly the South China Block. We are grateful to N. Arndt, B. Windley and T. Komiya for their constructive comments and suggestions. This study was financially supported by the Chinese Academy of Sciences (KZCX 1-107 and GIGCX-04-04) and the National Natural Science Foundation of China (40034343).

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Geochemistry of Mesozoic mafic volcanic rocks from the Yanshan belt in the northern margin of the North China Block: relations with post-collisional lithospheric extension F. GUO, W.-M. FAN, X.-Y. LI & C.-W. LI Key Laboratory of Marginal Sea Geology, Guangzhou Institute of Geochemistry, Chinese Academy of Sciences, Wushan, Guangzhou, 510640, China (e-mail: [email protected]) Abstract: Major element, trace element and Sr– Nd isotope data for three suites of mafic volcanic rocks that erupted at c. 180 Ma (Group 1), c. 163– 140 Ma (Group 2) and 136– 110 Ma (Group 3) in the Yanshan belt in the northern margin of the North China Block (NCB) are presented in this paper. All the rocks show significant enrichment in large ion lithophile elements and light REE (LREE) but depletion in Nb –Ta and Th–U, and moderately radiogenic Sr (87Sr/86Sr(i) ¼ 0.7052–0.7068) and unradiogenic Nd (1Nd(t) ¼ 215.1 to 27.2) isotopic compositions. Sr–Nd isotopic modelling suggests an insignificant role of crustal contamination or assimilation fractional crystallization, and the geochemical variations in these rocks were mainly attributed to source heterogeneity and variable degrees of ferromagnesian, plagioclase and accessory mineral fractionation. The low Th/La (0.039-0.10) points to recycling of low-Th/La (e.g. Th/La ,0.2) ancient crustal materials into the source region, probably related to the palaeo-Asian Ocean subduction. Compared with the late Mesozoic mafic rocks from the NCB interior, the Yanshan belt mafic lavas generally have higher Al2O3, TiO2, P2O5, LREE, high field strength elements, Sr and 87 Sr/86Sr(i) but lower MgO and compatible elements. Major element extrapolation (MgO ¼ 8% normalization) reveals that the Yanshan belt mafic volcanic rocks have higher Ti8 and Fe8, and lower Si8 than those from the NCB interior, suggesting that they were probably derived from a relatively fertile mantle source, different from the Mesozoic chemically refractory lithospheric mantle beneath the NCB interior. Combining the geochemical features of the mafic rocks with Mesozoic deformation events in the northern NCB, we suggest that the three stages of mafic volcanism were caused by episodic lithospheric extension. The Group 1 rocks, which occur locally along major faults, were generated during a post-compressional extension related to the collision between the NCB and Mongolian Block; the Group 2 rocks formed as a result of post-collisional lithospheric extension related to the collision between the North China–Mongolian Block and the Siberian plate; and the Group 3 rocks were extruded in an extensional regime in response to lateral escape related to surrounding plate interactions.

Mafic magmatism in orogenic belts provides important information on mantle–crust interaction, and chemical–thermal evolution of the lithosphere beneath plate margins (e.g. Menzies & Kyle 1990). The origin of such melts is commonly attributed to lithospheric extension by orogenic collapse (e.g. Ruppel 1995), slab breakoff (Davies & Brandenburg 1995), convective thinning (England & Houseman 1989) or delamination of continental lithosphere (Bird 1979; Kay et al. 1993). Unlike intraplate basalts (e.g. continental flood basalts, ocean island basalts (OIB) and mid-ocean ridge basalts (MORB), etc.), which usually show MORBor OIB-type geochemical characters, most orogenic mafic magmas exhibit arc-like trace element features (high large ion lithophile element to high field strength element ratios (LILE/HFSE) and light rare earth element (LREE)/HFSE values) and enriched isotopic signatures inherited from previous subduction- related enrichment (e.g. Hawkesworth et al. 1995; Turner et al. 1996; Jahn et al. 1999;

Miller et al. 1999; Fan et al. 2003, 2004; Guo et al. 2004). However, orogenic magmas are mainly derived from lithospheric mantle instead of a convective mantle wedge above a subducted slab (e.g. Hawkesworth et al. 1995; Turner et al. 1996; Fan et al. 2003). The geochemical features of these mafic rocks provide important constraints on geodynamic processes in lithospheric evolution. The North China Block (NCB) is known as one of the oldest continental blocks of the Earth, with crustal ages as old as c. 3.8 Ga (Liu et al. 1992). It is surrounded by the Triassic Qinling –Dabie – Sulu collisional orogen to the south and the late Palaeozoic to early Mesozoic NE China Fold Belt to the north (Fig. 1). The Yanshan belt is located between the NCB interior and the NE China Fold Belt (Fig. 1a; Sengor & Natal’in 1996; Davis et al. 2001). In the Mesozoic, episodic magmatism occurred coupled with multistage crust –lithosphere deformation events. The frequent tectonomagmatic events in this area have attracted the

From: ZHAI , M.-G., WINDLEY , B. F., KUSKY , T. M. & MENG , Q. R. (eds) Mesozoic Sub-Continental Lithospheric Thinning Under Eastern Asia. Geological Society, London, Special Publications, 280, 101– 130. DOI: 10.1144/SP280.5 0305-8719/07/$15 # The Geological Society of London 2007.

Fig. 1. (a) A simplified map showing the major tectonic units of the North China Block and its surrounding areas (modified from Davis et al. 2001). (b) A geological map showing the distribution of Mesozoic volcanic rocks in the Yanshan belt of the northern North China Block. CF-KY, Chifeng–Kaiyuan fault; CH-LP, Chicheng–Luanping fault; XL-KC, Xinglong–Kuancheng fault; LY-FN, Laiyuan–Fengning fault.

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interest of international geologists, and many hypotheses for their origin have been proposed; for example, the ‘Yanshan Movement’ (Wong 1929), intracontinental or intraplate orogeny (e.g. Zhang & Song 1997; Chen 1998; Zheng et al. 2000; Davis et al. 2001), and an active continental margin of the palaeo-Pacific Ocean (e.g. Deng et al. 1999; Faure & Natal’in 1992). To date, the major controversies on the Mesozoic lithospheric evolution are: (1) the role of the palaeo-Pacific plate in the accretion of the East Asian continental margin; (2) the relative contributions of the Mesozoic collision between the North China – Mongolian Block and the Siberian plate, and the collision between the North China and Yangtse Blocks; (3) the effect of plate interactions in the formation of the NE– SW trending structural lineation since the late Mesozoic. Although some studies are concerned with the Mesozoic magma origin in the northern NCB (e.g. Li et al. 2001; Shao et al. 2001; Zhou et al. 2001; Zhang et al. 2003), the nature of the subcontinental lithospheric mantle (SCLM), the relationship between structural deformation and magmatism, and the causes for the frequent tectonomagmatic events are still poorly understood. To further explore the geodynamics of Mesozoic tectono-magmatism in the Yanshan belt, we report here major element, trace element and Sr– Nd isotope data for three groups of Mesozoic mafic volcanic rocks, which were extruded in the early and late Jurassic, and the early Cretaceous. The purposes of this study are to: (1) constrain the nature of the Mesozoic SCLM beneath the northern NCB; (2) understand the genesis of these Mesozoic mafic rocks; (3) discuss the possible mechanisms for melt generation and its relationship with crust– lithosphere deformation.

Regional geology The Yanshan belt is the eastern segment of the east –west-trending Yinshan– Yanshan orogen in the northern NCB. The regional geology has been widely covered in previous studies (e.g. Bao et al. 1995; Zhai et al. 1996; Wu et al. 2000; Zhang & Song 1997; Davis et al. 2001; Shao et al. 2001; Cui et al. 2002; Liu et al. 2002). The Archaean to early Proterozoic was the most important crustal growth stage (e.g. Liu et al. 1992; Zhai et al. 1996). High-grade metamorphic terranes are widely exposed, intruded by widespread hypersthene granites and tonalite –trondhjemite–granites (TTG) of 2.4–2.5 Ga (e.g. Zhai et al. 1996). At c. 1.8 Ga, collision between the western and eastern blocks formed a north– south-trending high-pressure granulite belt (e.g. Zhao et al. 2000). The Archaean to

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early Proterozoic evolution resulted in the framework of the metamorphic basement. The Mesoproterozoic Yan-Liao passive rift is characterized by carbonates and terrestrial debris tens of kilometres thick (HBGMR 1989). The rift was closed by the end of the Mesoproterozoic. From the Neoproterozoic to early Palaeozoic, the Yanshan region became a passive continental margin of the palaeo-Asian Ocean, on which kilometre-scale neritic carbonates were deposited. Similar to the NCB interior, a stratigraphic gap from the Upper Ordovician to Lower Carboniferous marks uplift of the continent. During late Palaeozoic time, there developed a series of east– westtrending basins related to the closure of the palaeo-Asian Ocean, which contain predominant terrestrial deposits and extensive coal layers, and minor littoral sediments. The emplacement of the Indo-Sinian alkaline complexes may signify the termination of orogeny related to the palaeo-Asian Ocean (e.g. Yan et al. 2000). From early Mesozoic time, the Yanshan belt was remobilized and experienced many tectonomagmatic events associated with intracontinental deformation, thrust faults and folds, large-scale nappes and strike-slip faults, and metamorphic core complexes were well developed along this belt (e.g. Davis et al. 1998, 2001; He et al. 1998; Zheng et al. 2000; Wu et al. 2000). Subsequently, extensive volcanism, granitoid emplacement and terrestrial deposition tool place. Pre-180 Ma southvergent thrust faulting was prevalent in the northern NCB (e.g. Davis et al. 2001), followed by eruption of the Nandaling Formation (FM.) basalts. The second contractional deformation event took place at c. 175–163 Ma, at approximately by the time when the North China–Mongolian Block collided with the Siberian plate (e.g. Zhao et al. 1990; Davis et al. 2001). By the end of the Jurassic, there was an angular unconformity between the Zhangjiakou– Donglingtai and Houcheng –Tuchengzi Fms. In the early Cretaceous, regional deformation was dominated by extension, associated with extensive volcanism and formation of NE–SW trending fault-rift basins. Figure 2 summarizes the Mesozoic volcanosedimentary history of the Yanshan belt.

Petrography of Mesozoic mafic volcanic rocks Early Jurassic volcanic rocks (Group 1) are locally distributed along major faults; for example, the Xinglong–Kuancheng fault (also termed the Gubeikou– Pingquan fault; Davis et al. 2001) and the Chicheng –Luanping fault (Fig. 1b). The volcanic rocks, which mainly occur in the Xishan area near Beijing city and at the southern margin of the

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Fig. 2. A schematic summary of the Mesozoic volcano-sedimentary stratigraphic column for the Yanshan belt. The stratigraphic column for the Xishan area is compiled after BBGMR (1991), that for the Jibei area is after HBGMR (1989), and that for the Liaoxi area is after Chen et al. (1997). (See details in the text.)

Luanping basin in the Jibei region (Fig. 1b), crop out as subparallel layers or interbeds within the Lower Jurassic Nandaling Fm. The extrusive lavas truncate Upper Triassic shales, sandstones and conglomerates of the Xingshikou Fm., and are underlain unconformably by the Middle Jurassic Yaopo or Xiahuayuan Fm. They consist of basalts, basaltic

andesites and minor andesites that have K –Ar and Ar –Ar whole-rock eruption ages of 178–183 Ma (HBGMR 1989; Wang 1996; Chen 1998; Davis et al. 2001). Samples for this study were collected from the Xishan (20XSH-40 to -62) and Jibei (98LP-2 to -12) areas (Fig. 1b). The basalts are weakly porphyritic with predominant

GEOCHEMISTRY OF MESOZOIC MAFIC VOLCANIC ROCKS

phenocrysts of pyroxene and plagioclase, and minor olivine. The groundmass comprises fine-grained to aphanitic clinopyroxene and plagioclase, with minor ilmenite and other opaque oxides. Weak chloritization has occurred in basaltic samples, as reflected by their Na2O variation from 2.74 to 5.52% (Table 1). The basaltic andesites are also weakly porphyritic to aphanitic, with predominant clinopyroxene and plagioclase phenocrysts and subordinate amphibole. The matrix is mainly composed of fine-grained or aphanitic clinopyroxene and plagioclase (,0.2 mm), with minor opaque oxides. These rocks have also experienced slight alteration. The Group 2 rocks are widespread and comprise mainly intermediate –felsic volcanic rocks and minor basaltic lavas. The samples studied here were collected from the Tiaojishan and Houcheng Fms. Volcanic rocks from the Tiaojishan Fm. were erupted at 163– 147 Ma (Davis et al. 2001; Li et al. 2001; Niu et al. 2003). The basaltic lavas comprise mainly basaltic andesites and occur as interbeds in the intermediate –felsic rocks. The basaltic andesites are weakly porphyritic to aphanitic, with predominant clinopyroxene and plagioclase phenocrysts and subordinate amphibole, and show weak alteration. The matrix is mainly composed of fine-grained or aphanitic clinopyroxene and plagioclase (,0.2 mm), with minor opaque oxides. The Houcheng Fm. overlies the Tiaojishan Fm., and is overlain by rhyolites of the Zhangjiakou Fm., which has an oldest age of c. 136 Ma (Niu et al. 2003). The basaltic lavas that crop out as interbeds in syndepositional conglomerates are fairly fresh, subaphyric to weakly porphyritic, with predominant phenocrysts of clinopyroxene with rare olivine and plagioclase. The Group 3 lavas were extruded at the time of maximum volcanism in the early Cretaceous. They comprise predominant intermediate –felsic lavas and subordinate basalts, with an age range from 136 to 110 Ma (e.g. HBGMR 1989; Chen et al. 1997; Davis et al. 2001; Niu et al. 2003). The intermediate – felsic sequences in the Xishan and Jibei areas are termed the Donglintai and Zhangjiakou Fms., and the basaltic lavas are termed the Donglanggou and Huajiying Fms. (Fig. 2). The samples studied here were collected from both areas, and consist of trachybasalts and basaltic trachyandesites. The trachybasalts are porphyritic with predominant phenocrysts of pyroxene and plagioclase, with minor amphibole and rare olivine. The groundmass is composed of fine-grained clinopyroxene and plagioclase (,0.5 mm). The basaltic trachyandesites are weakly porphyritic with minor clinopyroxene, plagioclase and amphibole phenocrysts, and the matrix is composed of aphanitic materials and minor opaque oxides. Weak chloritization and

105

kaolinization are prevalent in both trachybasalts and basaltic trachyandesites.

Analytical techniques All samples were crushed to millimetre size after removal of weathered rims, and fresh chips were handpicked under a binocular microscope. The fresh chips were cleaned in an ultrasonic bath with deionized water. These chips were then crushed to ,20 mesh in a WC jaw crusher. A split was ground to ,160 mesh grain size in an agate ring mill for major and trace element analysis. Major elements were analysed at the Hubei Institute of Geology and Mineral Resource, Chinese Ministry of Land and Resources (MLR) by wavelengthdispersive X-ray fluorescence spectrometry (XRF) with analytical errors less than 2%. FeO content was analysed by a wet chemical method. The analysed standards have been reported by Fan et al. (2004). Trace element abundances of the samples were determined by inductively coupled plasma mass spectrometry (ICP-MS) at the Institute of Geochemistry, Chinese Academy of Sciences (CAS). The powders (about 100 mg) were placed in a screw-top PTFE-lined stainless steel bomb and dissolved by HF and HNO3. The sealed bombs were placed in an electric oven and heated to 190 8C for 12 h. Before ICP-MS analysis, 1 ppm Rh solution was added as an internal standard. The analytical errors are estimated to be less than 5% for most elements .10 ppm and about 10% for transition metals such as Cr, Ni, V and Sc, from repetitive analyses of the international standards BHVO-1 (basalt) and AMH-1 (andesite). Duplicate runs gave ,5% RSD (relative standard deviation) for most analysed elements, except for transition metals, which generally gave ,10% RSD. A detailed description of the analytical technique has been reported by Qi et al. (2000), and the analytical results for the standards BHVO-1 and AMH-1 are listed in Table 3. Sr and Nd isotopic ratios were determined by multi-collector inductively coupled plasma mass spectrometry (MC-ICP-MS) at Guangzhou Institute of Geochemistry, CAS. The Sr and Nd isotopic ratios were normalized to 86Sr/88Sr ¼ 0.1194 and 146 Nd/144Nd ¼ 0.7219, respectively. Ten analyses of the La Jolla standard gave 143Nd/144Nd ¼ 0.511862 + 10, and 12 analyses of BCR-1 gave 143 Nd/144Nd ¼ 0.512626 + 9. Seven analyses of NBS 987 yielded 87Sr/86Sr ¼ 0.710265 + 12. Total procedure blanks were about 2–5  10210g for Sr and less than 5  10211g for Nd. 87Rb/86Sr and 147Sm/144Nd ratios were calculated using Rb, Sr, Sm and Nd concentrations determined by ICP-MS. 87Sr/86Sr(i) and 1Nd(t) for the early

SiO2 TiO2 Al2O3 Fe2O3 FeO MnO MgO CaO Na2O K2O P2O5 LOI Total Sc V Cr Co Ni Ba Rb

49.24 1.92 15.58 3.01 7.22 0.15 5.62 7.73 3.11 1.63 0.76 3.64 99.61 24 200 113 39 61 978 24.1

49.24 2.11 15.37 1.88 8.00 0.16 6.44 6.44 3.82 1.47 0.81 3.88 99.62 25 202 119 35 52 1390 19.9

48.33 2.08 17.61 7.71 3.47 0.16 3.51 6.36 3.47 2.75 0.77 3.35 99.57 25 156 107 34 52 1585 46.6

48.40 2.14 16.48 6.49 5.70 0.16 4.71 5.11 4.14 1.57 0.81 3.92 99.63 25 177 119 47 71 950 26.8

47.64 2.03 16.98 3.96 6.85 0.16 5.17 6.28 4.12 1.51 0.81 4.13 99.64 24 174 132 39 70 932 22.1

53.52 54.07 52.98 1.00 1.15 0.70 18.9 18.02 19.37 2.61 3.45 6.13 4.49 4.15 1.97 0.12 0.11 0.08 3.10 3.80 2.40 4.80 7.10 4.50 4.23 3.10 5.07 1.11 0.67 1.08 0.60 0.77 0.60 4.70 3.10 3.98 99.40 99.89 99.31 14 15 14 115 142 148 39 76 47 26 27 23 25 30 32 2039 2540 730 14.9 5.8 22.9

52.33 0.72 19.13 6.72 0.18 0.11 3.40 4.75 5.43 1.11 0.64 5.00 99.82 13 113 49 19 27 900 10.9

54.06 1.10 17.52 4.01 3.30 0.09 4.63 3.31 6.05 1.69 0.36 3.54 99.66 17 159 40 25 22 1171 31.9

55.73 0.96 15.75 3.44 3.87 0.06 3.96 4.09 5.52 1.91 0.37 4.01 99.67 17 161 80 23 27 972 28.5

53.99 1.04 16.26 3.50 3.97 0.13 4.18 4.33 4.54 2.86 0.38 4.38 99.56 17 168 81 25 27 2113 77.9

52.41 1.24 16.84 3.67 4.47 0.08 5.75 2.73 4.29 2.60 0.49 5.09 99.66 17 167 16 24 12 1376 49.7

51.68 1.18 17.28 3.50 4.67 0.12 4.33 8.22 2.74 2.69 0.34 2.87 99.62 23 226 70 26 26 1212 52.3

Sample: 20XSH-40 20XSH-41 20XSH-42 20XSH-43 20XSH-44 98LP-2 98LP-4 98LP-8 98LP-12 20XSH-52 20XSH-57 20XSH-58 20XSH-59 20XSH-62 Location: Xishan Xishan Xishan Xishan Xishan Jibei Jibei Jibei Jibei Xishan Xishan Xishan Xishan Xishan Rock type: HTP-B HTP-B HTP-B HTP-B HTP-B LTP-BA LTP-BA LTP-BA LTP-BA LTP-BA LTP-BA LTP-BA LTP-BA LTP-BA

Table 1. Major (in wt%) and trace element (in ppm) compositions of representative Group 1 mafic volcanic rocks from the Yanshan belt

106 F. GUO ET AL.

1243 2.30 0.67 274 6.43 17.2 0.78 35.7 46.85 104.5 12.36 51.81 9.40 2.64 8.15 1.21 6.57 1.27 3.45 0.51 3.21 0.44

541 2.40 0.64 286 6.96 18.7 0.91 38.5 49.08 109.1 12.49 52.16 9.84 2.47 8.43 1.27 7.27 1.35 3.80 0.56 3.38 0.48

1173 2.54 0.71 294 7.23 18.4 0.91 35.2 50.48 106.4 12.85 51.26 9.22 2.58 7.91 1.21 6.65 1.26 3.69 0.54 3.22 0.44

702 2.70 0.70 310 7.60 19.8 0.96 39.3 54.54 118.5 14.09 57.78 10.77 2.83 8.95 1.32 7.53 1.46 3.92 0.57 3.62 0.54

966 2.49 0.59 290 6.73 18.3 0.84 38.1 50.54 109.9 13.16 55.35 10.61 2.72 8.79 1.24 7.06 1.35 3.66 0.52 3.34 0.44

HTP: high-Ti-P series; LTP: low-Ti-P series; B, basalt; BA, basaltic andesite.

Sr Th U Zr Hf Nb Ta Y La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu

1313 1400 998 3.41 2.61 3.50 0.83 0.58 0.89 215 202 226 5.10 5.20 5.40 9.3 9.3 9.9 0.47 0.53 0.53 18.0 18.5 18.8 42.47 39.80 45.44 91.3 84.3 96.7 10.73 9.69 11.45 43.49 39.74 45.70 7.97 7.79 8.08 1.95 1.86 2.05 5.84 5.21 5.74 0.78 0.74 0.77 3.67 3.69 3.94 0.70 0.70 0.76 1.92 1.98 1.87 0.25 0.23 0.24 1.51 1.67 1.69 0.23 0.27 0.21

988 3.67 1.18 220 5.13 9.9 0.57 18.2 43.91 93.1 10.41 42.40 7.87 2.04 5.63 0.72 3.87 0.70 1.88 0.26 1.48 0.20

750 2.44 0.51 155 4.23 7.8 0.36 16.0 33.28 70.0 8.24 34.37 6.16 1.66 4.50 0.59 3.17 0.51 1.41 0.20 1.17 0.18

708 2.87 0.61 165 4.39 9.5 0.41 15.8 37.24 76.7 8.93 34.81 6.02 1.58 4.47 0.59 2.82 0.51 1.45 0.20 1.26 0.16

958 3.05 0.66 170 4.48 9.8 0.42 16.1 38.68 80.2 9.24 35.97 6.19 1.69 4.71 0.59 3.03 0.53 1.51 0.21 1.31 0.17

668 2.65 0.74 198 5.46 10.6 0.48 25.9 34.01 80.1 10.28 42.99 8.08 2.09 6.60 0.84 4.86 0.91 2.35 0.34 2.10 0.29

999 2.06 0.49 121 3.65 6.5 0.30 18.3 23.42 52.9 6.62 28.05 5.56 1.65 4.41 0.65 3.54 0.65 1.85 0.25 1.62 0.22

GEOCHEMISTRY OF MESOZOIC MAFIC VOLCANIC ROCKS 107

SiO2 TiO2 Al2O3 Fe2O3 FeO MnO MgO CaO Na2O K2O P2O5 LOI Total Sc V Cr Co Ni Ba Rb Sr Th U

49.34 1.79 17.34 9.09 1.75 0.18 3.04 6.54 4.31 2.70 0.82 2.53 99.43 22 270 13 33 14 2033 43.2 1929 3.68 0.73

49.36 1.80 17.27 6.22 3.33 0.18 4.45 6.44 3.86 2.95 0.84 2.80 99.50 21 249 12 32 14 1846 53.3 1557 3.69 0.72

49.75 1.76 17.19 9.13 1.83 0.15 3.42 5.90 4.58 2.60 0.82 2.38 99.51 22 259 14 32 15 1611 48.1 1862 3.67 0.80

50.91 1.90 15.50 8.90 1.47 0.10 3.79 6.79 3.63 2.74 0.99 2.78 99.50 20 185 99 30 42 1923 48.1 1454 3.87 0.81

51.88 1.87 15.35 8.14 1.73 0.11 3.00 6.57 4.22 2.97 0.96 2.75 99.55 19 218 98 27 39 1581 58.0 1361 3.70 0.78

51.39 53.27 55.12 56.03 51.25 1.68 1.15 1.00 1.05 1.02 15.01 19.13 19.06 18.02 19.96 8.25 5.81 5.91 6.06 3.68 1.90 2.09 0.99 1.24 4.93 0.13 0.09 0.09 0.13 0.10 4.06 2.70 2.30 2.50 3.29 6.15 4.80 4.60 4.90 5.55 3.98 4.38 4.66 4.15 4.20 3.13 2.44 2.22 2.48 2.00 0.91 0.70 0.93 0.86 0.44 2.94 3.00 2.70 2.10 3.17 99.53 99.56 99.58 99.52 99.59 18 11 12 11 12 155 139 152 138 165 115 50 51 37 23 32 92 43 22 23 55 26 27 24 14 1649 1268 1248 1312 1482 61.9 42.7 40.4 44.6 26.8 1345 1012 972 845 1242 3.68 1.54 1.50 1.50 2.25 0.73 0.43 0.53 0.44 0.55

52.28 1.13 17.61 5.69 1.27 0.13 3.23 7.65 4.48 1.69 0.5 3.95 99.61 14 157 51 23 23 1175 27.3 1408 1.45 0.50

56.97 1.10 16.77 5.45 1.47 0.09 3.51 4.04 5.58 1.50 0.48 2.67 99.63 13 132 47 22 20 1073 28.8 1263 1.41 0.56

57.43 0.88 15.34 2.86 3.85 0.07 3.60 6.22 3.26 2.63 0.32 3.17 99.63 15 147 42 20 15 1115 69.0 1138 3.37 0.72

56.99 1.03 15.41 5.44 1.93 0.08 3.90 5.81 4.16 1.73 0.33 2.83 99.64 15 166 18 22 18 1204 28.5 970 2.46 0.45

Sample 20XSH-12220XSH-12320XSH-12520XSH-13020XSH-13220XSH-13398LP-2498LP-2698LP-2720XSH-6820XSH-7620XSH-7720XSH-8520XSH-88 Location Xishan Xishan Xishan Xishan Xishan Xishan Jibei Jibei Jibei Xishan Xishan Xishan Xishan Xishan Rock TB-H TB-H TB-H TB-H TBA-H TBA-H BA-T BA-T A-T BA-T BA-T A-T A-T A-T type*

Table 2. Major (in wt%) and trace element (in ppm) compositions of representative Group 2 mafic volcanic rocks from the Yanshan belt

108 F. GUO ET AL.

267 6.45 20.9 0.97 29.1 54.88 121.6 14.33 57.11 9.76 2.64 7.81 1.06 5.56 1.01 2.64 0.39 2.39 0.35

260 6.23 21.0 0.95 28.2 55.58 119.8 14.22 57.47 10.06 2.60 7.62 1.06 5.38 1.02 2.66 0.40 2.42 0.32

265 6.43 21.0 0.95 28.9 55.49 120.4 14.24 56.56 10.12 2.64 7.53 1.01 5.37 1.03 2.78 0.39 2.38 0.33

349 8.02 25.1 1.03 27.7 77.05 159.0 18.42 71.70 11.55 2.75 8.22 1.07 5.38 0.94 2.61 0.35 2.12 0.31

370 8.53 27.0 1.14 29.0 80.14 167.5 19.44 74.73 12.22 2.79 8.80 1.12 5.86 1.03 2.79 0.36 2.23 0.29

355 8.15 25.8 1.06 27.9 78.17 162.7 18.57 72.31 11.83 2.90 8.43 1.09 5.37 0.99 2.69 0.38 2.30 0.29

TB, trachybasalt; BTA, basaltic trachyandesite; A, andesite; H, Houcheng Fm.; T; Tiaojishan Fm.

Zr Hf Nb Ta Y La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu

376 186 183 8.65 4.57 4.69 25.0 7.6 7.6 1.04 0.43 0.38 27.0 13.5 14.0 80.40 33.17 33.53 163.9 74.80 74.64 18.63 9.06 9.17 71.04 39.05 39.76 11.63 7.72 7.64 2.68 1.92 1.93 7.99 5.10 5.26 1.09 0.62 0.65 5.15 3.28 3.13 0.93 0.51 0.48 2.50 1.38 1.35 0.35 0.15 0.13 2.12 1.01 1.09 0.30 0.13 0.14

180 4.71 7.7 0.36 13.1 32.39 73.23 9.03 37.77 7.18 1.78 4.59 0.59 2.86 0.46 1.21 0.15 0.99 0.15

172 5.30 8.6 0.35 17.9 34.67 75.23 9.02 36.10 6.50 1.75 4.95 0.60 3.21 0.58 1.57 0.23 1.41 0.19

162 4.34 8.2 0.32 14.8 35.40 76.75 9.48 38.96 6.91 1.89 4.97 0.59 2.88 0.47 1.32 0.17 1.04 0.14

153 4.22 8.0 0.33 13.7 33.85 74.70 9.10 37.22 6.59 1.78 4.60 0.56 2.74 0.44 1.17 0.15 1.00 0.13

154 150 4.33 3.78 8.4 7.9 0.54 0.34 15.3 14.8 33.75 30.52 69.29 65.18 7.94 7.83 31.18 31.17 5.54 5.49 1.45 1.49 4.12 4.15 0.54 0.55 2.86 2.83 0.54 0.49 1.37 1.33 0.20 0.19 1.32 1.19 0.19 0.16

GEOCHEMISTRY OF MESOZOIC MAFIC VOLCANIC ROCKS 109

51.76 1.85 19.73 2.92 5.38 0.13 2.50 6.00 4.08 1.49 1.33 2.18 99.85 12 111 46 31 21 1136 38.1 749 4.94 1.08

SiO2 TiO2 Al2O3 Fe2O3 FeO MnO MgO CaO Na2O K2O P2O5 LOI Total Sc V Cr Co Ni Ba Rb Sr Th U

52.57 2.00 16.53 3.59 5.71 0.14 3.20 6.20 3.75 1.53 1.50 2.38 99.80 15 147 57 164 30 1290 60.0 916 4.81 1.04

BTA-H

Rock type: BTA-H

51.97 2.00 17.01 3.53 5.47 0.09 3.10 5.60 3.55 2.46 1.47 2.70 99.37 15 152 48 64 24 1236 67.0 800 6.09 1.25

BTA-H

TB-H

BTA-H

BTA-H

49.41 50.87 53.54 54.00 1.85 2.20 1.38 1.37 18.66 17.48 18.10 17.71 3.05 4.41 2.90 3.98 5.25 5.69 5.2 5.02 0.14 0.14 0.14 0.15 2.30 3.20 2.50 2.40 7.20 5.80 4.9 5.10 3.29 3.57 4.14 3.92 2.47 2.26 2.31 2.73 1.50 1.37 0.97 1.27 2.50 1.90 3.07 1.70 99.62 99.29 99.75 99.65 16 14 12 14 155 155 112 119 84 54 32 40 27 35 22 35 24 29 18 19 1054 1431 1381 1331 71.6 99.7 123.9 57.8 771 938 848 842 6.25 5.17 4.46 4.12 1.34 1.16 0.91 0.91

TB-H 49.59 1.86 15.94 4.63 4.33 0.17 5.37 7.56 2.33 3.72 1.02 2.81 99.33 18 181 148 33 73 3849 49.2 1152 3.12 0.80

TA-D 53.72 1.54 15.45 5.28 2.32 0.11 3.57 7.19 3.38 3.40 0.93 2.55 99.44 14 147 122 25 64 2660 46.1 1293 2.91 0.70

BTA-D 48.36 1.99 15.79 4.48 4.22 0.19 5.72 8.30 2.38 3.83 1.04 2.95 99.25 18 187 156 31 67 4743 47.0 1194 2.91 0.76

TB-D

54.55 1.60 15.66 6.10 1.73 0.06 2.91 5.75 3.34 4.70 0.99 1.90 99.29 15 160 126 22 56 3618 63.6 1415 3.14 0.70

BTA-D

53.47 1.60 15.73 6.41 1.93 0.07 2.8 6.59 3.22 4.32 0.98 2.23 99.35 15 163 127 24 63 3149 64.2 1512 3.17 0.70

BTA-D

Sample: 98WC-15 98WC-16 98WC-19 98WC-20 98FN-14 98FN-18 98FN-20 20XSH-29 20XSH-30 20XSH-31 20XSH-32 20XSH-33 Location: Jibei Jibei Jibei Jibei Jibei Jibei Jibei Xishan Xishan Xishan Xishan Xishan

Table 3. Major (in wt%) and trace element (in ppm) compositions of representative Group 3 mafic volcanic rocks from the Yanshan belt

13.9 + 0.2 114 + 3 41 + 3 19.5 + 0.3 35.5 + 0.6 315 + 10 18.3 + 0.2 562 + 4 2.61 + 0.02 0.91 + 0.03

AMH-1 This work (n ¼ 6) andesite

31.7 + 0.4 318 + 4 275 + 12 44.7 + 0.4 121 + 2 138 + 5 9.8 + 0.5 396 + 10 1.27 + 0.04 0.43 + 0.02

BHVO-1 This work (n ¼ 15) basalt

110 F. GUO ET AL.

431 9.01 26.9 1.50 29.1 75.33 160.5 18.93 76.74 14.90 3.14 9.76 1.27 6.58 1.10 2.77 0.35 2.19 0.32

396 8.29 21.3 1.29 22.9 67.07 141.7 15.84 62.21 11.54 2.30 7.64 1.00 5.25 0.84 2.26 0.28 1.87 0.27

418 9.19 28.8 1.44 26.4 78.19 162.1 18.87 72.77 13.21 2.98 9.23 1.16 6.03 1.01 2.57 0.29 2.00 0.28

434 9.44 28.8 1.50 24.8 79.19 166.6 18.70 76.06 13.61 2.94 9.13 1.02 5.77 0.97 2.50 0.30 1.87 0.29

466 10.01 31.5 1.63 30.1 86.29 179.7 20.73 83.27 15.67 3.36 10.38 1.30 6.83 1.18 3.22 0.39 2.37 0.33

437 9.82 31.7 1.62 32.7 66.98 141.8 16.03 62.65 12.00 2.72 9.00 1.23 6.85 1.25 3.47 0.45 3.06 0.41

Huajiying Fm.; D, Donglanggou Fm.; other abbreviations are the same as in Tables 1 and 2.

Zr Hf Nb Ta Y La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu

394 8.96 35.0 1.86 32.8 66.41 142.6 16.53 65.80 12.07 2.81 9.71 1.29 7.34 1.31 3.46 0.45 3.13 0.41

245 6.13 29.1 1.39 21.1 69.98 150.7 17.35 70.89 11.00 2.63 7.54 0.95 4.36 0.75 1.86 0.23 1.60 0.20

251 6.30 23.4 1.05 17.7 75.08 159.0 18.75 72.62 10.87 2.59 7.11 0.82 3.65 0.59 1.59 0.21 1.35 0.15

254 6.56 30.6 1.51 20.6 72.97 158.7 18.95 72.72 11.80 2.83 7.79 0.93 4.47 0.74 1.90 0.24 1.50 0.19

271 6.80 25.1 1.15 18.8 81.24 172.1 20.08 77.39 11.60 2.82 7.53 0.87 4.03 0.66 1.74 0.24 1.31 0.18

268 139 + 5 6.74 3.8 + 0.1 25.1 8.4 + 0.2 1.15 0.60 + 0.1 19.0 15.2 + 0.2 81.46 16.22 + 0.15 171.9 34.5 + 0.3 19.93 4.2 + 0.2 77.62 17.9 + 0.2 11.47 3.8 + 0.1 2.64 1.08 + 0.05 7.66 3.48 + 0.06 0.84 0.52 + 0.01 4.08 3.01 + 0.03 0.67 0.58 + 0.10 1.65 1.54 + 0.03 0.22 0.22 + 0.02 1.52 1.35 + 0.05 0.18 0.22 + 0.01

176 + 6 4.6 + 0.2 19.2 + 0.2 1.18 + 0.02 28.0 + 0.2 15.16 + 0.38 37.13 + 0.84 5.6 + 0.2 25.9 + 0.8 6.3 + 0.2 2.04 + 0.05 6.35 + 0.30 0.99 + 0.04 5.20 + 0.11 0.96 + 0.02 2.37 + 0.02 0.34 + 0.01 2.00 + 0.02 0.30 + 0.02

GEOCHEMISTRY OF MESOZOIC MAFIC VOLCANIC ROCKS 111

112

F. GUO ET AL.

Jurassic, late Jurassic and early Cretaceous basalts were calculated using ages of 180 Ma, 150 Ma and 120 Ma. Major oxide and trace element compositions and Sr–Nd isotopic data for the samples are listed in Tables 2 –4.

series: (1) high-Ti and high-P (termed J1-HTP) and (2) low-Ti and low-P (termed J1-LTP). The J1-HTP series have lower SiO2 and Al2O3 but higher HFSE and REE concentrations than the J1-LTP series. There are no systematic elemental correlations between the two series, implying that they were not comagmatic (Li et al. 2004).

Results Major and trace elements On a loss on ignition (LOI)-free basis, the studied mafic rocks have 49.9– 59.5% SiO2, 2.4–6.7% MgO, and a Mg-number (Mg-number ¼ 100  Mg/(Mg þ FeT) in atomic ratio) range of 34 –57. They have medium-K calc-alkaline to high-K calc-alkaline, or even shoshonitic, affinities (Fig. 3). In essence, the Mesozoic Yanshan belt mafic rocks generally have low MgO (or Mg-number), and high Al2O3, TiO2, P2O5, Sr, LREE and HFSE relative to the late Mesozoic mafic rocks from the NCB interior (Guo et al. 2001, 2003; Fig. 4); these are features of differentiated melts. The main elemental variations of the three basaltic suites are summarized below. Group 1. The Group 1 rocks have a wide variation in SiO2 and K2O (Fig. 3), showing medium-K to high-K calc-alkaline affinities. They span large variations in TiO2 and P2O5 coupled with change in other major and trace elements. On the basis of the compositional gap in TiO2 and P2O5 (Table 1 and Fig. 4), we divided these rocks into two

Fig. 3. SiO2 v. K2O classification diagram for the Mesozoic Yanshan belt basalts (after Middlemost 1994). The late Mesozoic NCB interior mafic rocks (Guo et al. 2001, 2003, and unpublished data) are plotted for comparison (the shaded area). J1-HTP, high-Ti – P series from the Nandaling Fm.; J1-LTP, low-Ti–P series from the Nandaling Fm.; J3-T, samples from the Tiaojishan Fm.; J3-H, samples from the Houcheng Fm.; K1-D, samples from the Donglingtai Fm.; K1-H, samples from the Huajiying Fm.

Group 2. Like Group 1, the Group 2 rocks are also composed of a low-Ti –P series in the Tiaojishan Fm. (termed J3-T) and a high-Ti–P series in the Houcheng Fm. (termed J3-H). The J3-T samples have higher Al2O3 (17.5–20.0%) and Sr (845 – 1408 ppm), low Y (13.1– 17.9 ppm), high Sr/Y (52– 95) and strong LREE/HREE (heavy REE) fractionation (i.e. La/YbCN ¼ 15.9–23.0), features of adakitic magmas (Fig. 5; e.g. Defant & Drummond, 1990; Martin et al. 2005). A slight change in REE, Th and Nb, and a decrease in FeOT, Al2O3, CaO and Sr are coincident with an increase in SiO2. The J3-H basalts show shoshonitic affinities and narrow changes in major and trace elements (e.g. SiO2 49.4 –52.1%, MgO 2.6–4.5%) except for La (54.9–88.4 ppm) and Sr (1317–2033 ppm). They are evolved magmas, according to their low MgO and Cr (12–115 ppm) and Ni (16–24 ppm) concentrations. Group 3. The Group 3 mafic volcanic rocks span a SiO2 range of 50.2–56.0% and a MgO range of 2.4–5.9%, showing high-K calc-alkaline to shoshonitic affinities (Fig. 3). They generally have higher LREE and HFSE concentrations than the other two groups (Table 3 and Fig. 4). Samples from the Huajiying Fm. (termed K1-H) have the lowest MgO and highest P2O5 and Th of the studied rocks (Table 3). Following an increase in SiO2, they show decrease in FeOT, CaO, TiO2, P2O5, LREE and Th, an increase in Nb, and weak Sr variation (Fig. 4). The other rocks from the Donglingtai Fm. (termed K1-D) demonstrate a decrease in MgO, FeOT, TiO2, Nb and Sr, an increase in LREE, and weak variation in Al2O3, P2O5 and Th following magma evolution. All the Mesozoic Yanshan belt mafic volcanic rocks are characterized by significant LILE and LREE enrichment, and depletion in Nb–Ta (e.g. La/Nb ¼ 1.9–4.6 and Ba/Nb ¼ 37– 273) and Th –U (e.g. Th/La ¼ 0.039–0.10, with an average of 0.062) in primitive mantle-normalized incompatible element spidergrams (Fig. 6), features very different from those of MORB and OIB (Sun & McDonough 1989).

Sr and Nd isotopes The analytical and age-corrected Sr and Nd isotopic data for the Mesozoic Yanshan belt mafic lavas are

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Fig. 4. SiO2 v. major elements (in wt%) and trace elements (in ppm) for the Mesozoic Yanshan belt mafic rocks. The shaded areas show the variation fields of the late Mesozoic NCB interior mafic rocks. Data sources and symbols are as in Figure 3.

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Fig. 5. Sr v. Sr/Y for the Mesozoic Yanshan belt mafic volcanic rocks. It should be noted that some J3-H and J1-LTP samples have low Y and high Sr/Y ratios similar to modern adakites or Archaean tonalite– trondhjemite– granodiorites.

listed in Table 4. These rocks generally exhibit fairly large variations in Sr and Nd isotope compositions. In Group 1, two J1-HTP samples have initial 87 Sr/86Sr(i) of c. 0.7060 and 1Nd(t) of c. 27.5, similar to the high-Mg andesites from the northern NCB (Zhang et al. 2003), but different from the J1-HTP rocks, which have an 87Sr/86Sr(i) range of 0.7052–0.7061 and 1Nd(t) range from 215.1 to 212.1. In Group 2, the J3-H samples span an 87 Sr/86Sr(i) range of 0.7053–0.7061 and an 1Nd(t) range from 214.9 to 213.5, overlapping the variation field of the J1-HTP rocks (Fig. 7); 87 Sr/86Sr(i) ranges from 0.7064 to 0.7066 and 1Nd(t) from 214.1 to 211.8 for the J3-H rocks. In Group 3, the K1-D rocks have an 87Sr/86Sr(i) range of 0.7057–0.7059, 1Nd(t) from 214.8 to 212.5, and less radiogenic Sr and Nd isotopic compositions than the K1-H samples, which span an

Fig. 6. Primitive mantle-normalized incompatible element spidergrams of Mesozoic Yanshan belt basalts. All the basaltic samples show remarkable Th– U and Nb –Ta negative anomalies relative to Ba and La. Data for the late Mesozoic NCB interior mafic rocks are from Guo et al. (2001, 2003); data for PM, OIB and N-MORB are from (Sun & McDonough 1989).

1313 1400 998 1243 541 708 958 668

1012 972 845 1263 1138 970 1929 1557 1862 1361

916 800 771 938 842 1152 1293 1194 1415

42.7 40.4 44.6 28.8 67.0 28.5 43.2 53.3 48.1 58.0

60.0 67.0 71.6 99.7 57.8 49.2 46.1 47.0 63.6

Sr (ppm)

14.9 5.76 22.9 24.1 19.9 28.5 77.9 49.7

Rb (ppm)

14.90 13.21 13.61 15.67 12.07 11.00 10.87 11.80 11.60

7.72 7.64 7.18 6.59 5.54 5.49 9.76 10.06 10.12 11.83

7.97 7.79 8.08 9.40 9.84 6.02 6.19 8.08

Sm (ppm)

76.74 72.77 76.06 83.27 65.80 70.89 72.62 72.72 77.39

39.0 39.8 37.8 37.22 31.18 31.17 57.11 57.47 56.56 72.31

43.49 39.74 45.70 51.81 52.16 34.81 35.97 42.99

Nd (ppm)

0.1894 0.1471 0.2684 0.3077 0.1986 0.1237 0.1031 0.1138 0.1300

0.1220 0.1204 0.1528 0.0661 0.1756 0.0852 0.0649 0.0991 0.0748 0.1223

0.0328 0.0119 0.0664 0.0562 0.1067 0.1163 0.2354 0.2154

Rb/86Sr

87

0.7053 0.7054 0.7054 0.7061 0.7058 0.7058 0.7066 0.7066 0.7066 0.7064 0.7065 0.7068 0.7065 0.7067 0.7063 0.7058 0.7058 0.7057 0.7059

0.705547 + 17 0.705672 + 17 0.705651 + 18 0.706133 + 16 0.706183 + 11 0.705937 + 10 0.706766 + 10 0.706818 + 10 0.706931 + 20 0.706679 + 18 0.706861 + 14 0.707067 + 16 0.707002 + 18 0.707216 + 13 0.706623 + 17 0.705984 + 16 0.705939 + 14 0.705901 + 16 0.706087 + 18

Sr/86Sr(i)* 0.7052 0.7054 0.7055 0.7059 0.7061 0.7061 0.7060 0.7059

87

0.705308 + 14 0.705443 + 16 0.705631 + 16 0.706091 + 14 0.706346 + 16 0.706383 + 13 0.706629 + 17 0.706404 + 17

Sr/86Sr + 2s

87

*The initial Sr and Nd isotope data Group 1, 2 and 3 basalts are recalculated using a mean age of 180 Ma, 150 Ma and 120 Ma, respectively.

Group 1 98LP-2 98LP-4 98LP-8 20XSH-40 20XSH-41 20XSH-57 20XSH-58 20XSH-59 Group 2 98LP-24 98LP-26 98LP-27 20XSH-77 20XSH-85 20XSH-88 20XSH-122 20XSH-123 20XSH-125 20XSH-132 Group 3 98WC-15 98WC-19 98WC-20 98FN-14 98FN-20 20XSH-29 20XSH-30 20XSH-31 20XSH-32

Sample

Table 4. Sr and Nd isotope data of Mesozoic mafic volcanic rocks from the Yanshan belt 147

0.1174 0.1097 0.1081 0.1138 0.1109 0.0938 0.0905 0.0981 0.0906

0.1196 0.1162 0.1149 0.1070 0.1074 0.1065 0.1033 0.1058 0.0976 0.0989

0.1107 0.1185 0.1069 0.1097 0.1141 0.1045 0.1040 0.1137

Sm/144Nd

Nd/144Nd + 2s

0.512114 + 9 0.512036 + 8 0.512046 + 8 0.511997 + 8 0.512007 + 10 0.511855 + 8 0.511735 + 8 0.511847 + 8 0.511805 + 11

0.511819 + 9 0.511832 + 8 0.511797 + 12 0.511806 + 11 0.511830 + 11 0.511856 + 12 0.511856 + 11 0.511853 + 10 0.511817 + 9 0.511733 + 10

0.511873 + 10 0.511632 + 6 0.511721 + 10 0.512020 + 11 0.512018 + 13 0.511773 + 13 0.511784 + 12 0.511711 + 11

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27.2 28.7 28.5 29.5 29.3 212.5 214.8 212.7 213.5

214.6 214.4 214.9 214.5 214.1 213.5 211.8 211.8 212.5 214.1

212.9 215.1 213.4 27.5 27.6 212.4 212.1 213.6

1Nd(t)

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Fig. 7. 87Sr/86Sr(i) v. 1Nd(t) for the Mesozoic Yanshan belt basalts. Data for the high-Mg andesites are from Zhang et al. (2003); for the Triassic alkaline intrusive complexes from Yan et al. (2000); for the NCB upper crust from our unpublished data for Cenozoic NCB sediments); for granulite-facies xenoliths from Zhang et al. (1998); Zhou et al. (2002); Huang et al. (2004) and Li Huang et al. (2004); for the late Mesozoic NCB interior mafic rocks from Guo et al. (2001, 2003, and unpublished data). 87

Sr/86Sr(i) range of 0.7063–0.7068 and 1Nd(t) range of 29.5 to 27.2. In essence, the Yanshan belt mafic rocks generally have higher 87Sr/86Sr(i) ratios (0.7052– 0.7068) than those of late Mesozoic counterparts from the NCB interior (Fig. 7), which span an 87 Sr/86Sr(i) range of 0.7040– 0.7057 (Guo et al. 2001, 2003).

crystallisation (AFC). Below, we will evaluate the effect of both magmatic processes and source characteristics on magma evolution. In addition, a geochemical comparison with the late Mesozoic NCB interior mafic rocks is also presented, to show the possible difference of SCLM between cratonic and circumcratonic regions.

Crustal contamination Petrogenesis of the Mesozoic Yanshan belt mafic lavas The Mesozoic Yanshan belt mafic lavas have MgO, Cr and Ni values too low to represent primary mantle-derived magmas (e.g. Hart & Davis 1978). To account for their enriched character (e.g. LREE and LILE and Sr–Nd isotopes), the petrogenesis of the three groups necessities the involvement of: (1) an enriched component, either asthenospheric (e.g. a deep mantle plume) or lithospheric in origin, or (2) enrichment processes, such as high degrees of fractional crystallization, crustal contamination or assimilation –fractional

Basalts erupted in continental settings are often enriched in incompatible trace elements and isotopic features relative to their oceanic counterparts. The debate continues as to the importance of processes operating at crustal levels, which makes it necessary to evaluate the possible effects of crustal contamination or AFC. All the Mesozoic Yanshan belt mafic rocks show significant Nb–Ta depletion relative to LILE and LREE, radiogenic Sr and unradiogenic Nd isotopic compositions. Mass balance consideration suggests that such geochemical features in the Yanshan belt mafic rocks cannot be realistically attributed to crustal contamination of asthenosphere-derived melts

GEOCHEMISTRY OF MESOZOIC MAFIC VOLCANIC ROCKS

such as normal MORB (N-MORB) and OIB (e.g. Hawkesworth et al. 1995; Jahn et al. 1999; Fan et al. 2003). However, if the primary magmas were derived from enriched mantle sources, the variable Sr–Nd isotope data for the three groups might have been a result of crustal contamination or AFC. To discuss the role of crustal contamination and/or AFC during magmatic evolution, we selected the possible endmember components to model these possibilities. Sample 98LP-2 has the lowest 87Sr/86Sr(i), and thus it is assumed to represent the least contaminated primary melt for Sr isotopic modelling. Because we have only two HTP samples from Group 1, it is difficult to evaluate the role of crustal contamination in their petrogenesis, but the fact that they have the highest MgO and almost the highest 1Nd(t) of all studied rocks indicates that they represent the least contaminated melts. Here we select sample 20XSH-41, which has the highest MgO (6.44%) and higher 1Nd(t) value (27.6), to represent the least contaminated magma for Nd isotopic calculation. The modelling results for both crustal contamination and AFC processes are illustrated in Figure 8. The K1-H samples, which have the lowest MgO and the highest 87Sr/86Sr(i) of all the studied rocks, are the most likely candidates for contaminated melts. Sr isotopic modelling results require addition of 40 –50% NCB upper continental const (UCC) to magma like 98LP-2 to produce the higher 87 Sr/86Sr(i) in these rocks (Fig. 8a). However, addition of such a large volume of cold NCB UCC could not be re-equilibrated for heat budget. AFC processes could be another possibility to generate their higher 87Sr/86Sr(i). The AFC modelling result shows that with a KSr D range of 0.8– 1.2 and remaining magma fraction (F) ,0.3 (Fig. 8b), these rocks could be produced by assimilation of the NCB UCC. However, such a low F would lead to a high extent of ferromagnesian fractionation during magma evolution, which should have created an evolved magma of felsic rather than mafic composition as observed at present. Accordingly, neither crustal contamination nor AFC was predominant in the magma evolution of the K1-H rocks. The J1-LTP and J3-T samples, which have 1Nd(t) 4–8 units lower than the J1-HTP and K1-H rocks, may have resulted from crustal contamination and/or AFC. As illustrated in Figure 8c, both the J1-LTP and J3-T samples could be formed by addition of 50–80% NCB UCC or 40– 60% NCB lower or middle continental crust (LCC or MCC) to magma like 20XSH-41. For heat budget reasons, such a high volume of contamination is unlikely to happen. AFC could be another process to produce their lower 1Nd(t). The AFC modelling

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result indicates that with a KNd D range of 0.8–1.2 and remaining magma fraction F ,0.2, assimilation of either the NCB UCC or LCC –MCC to melt like 20XSH-41 would produce the Nd isotopic compositions in these two groups of rocks (Fig. 8d). Similarly, the magma produced should have been felsic rather than mafic composition. Therefore crustal contamination or AFC should be subordinate even if it did occur. The remaining J3-H and K1-D rocks, which plot away from the crustal contamination trend and have less varied Sr and Nd isotope data compared with the other groups from the Yanshan belt, escaped significant crustal contamination. For instance, Sr isotopic modelling results show that with F ,0.4, assimilation of either NCB UCC or NCB LCC – MCC would create the observed Sr isotopic compositions in these samples (Fig. 8b); whereas with F ,0.2, assimilation of either NCB UCC or NCB LCC –MCC would generate the observed Nd isotopic compositions (Fig. 8d). In summary, crustal contamination or AFC had an insignificant effect on the magma evolution of the studied Mesozoic Yanshan belt mafic rocks. The Sr–Nd isotopic variations in these rocks were mainly inherited from a heterogeneous source.

Fractional crystallization As mentioned above, all the studied mafic lavas from the Yanshan belt show an evolved affinity, with low MgO and compatible trace element contents, which would require a significant role of fractional crystallization during magma evolution. In Group 1, the compositional gaps in TiO2, P2O5 and other incompatible elements (e.g. REE and HFSE) between the J1-HTP and J1-LTP series cannot be produced through fractional crystallization of a common primary magma (Li et al. 2004), so the two series had distinct fractional phases during their passage through the lithosphere. The J1-HTP samples show weak variations in major and trace elements and have the highest MgO, indicating an insignificant role of fractional crystallization en route to the surface. In contrast, fractionation of a few percent of ferromagnesian phases and plagioclase and minor Ti-rich phases (e.g. rutile, ilmenite and Ti-magnetite) is necessary to account for the broad decrease in MgO, FeOT, CaO, Al2O3, TiO2 and HFSE in the J1-LTP series. Similarly, the compositional gaps in Al2O3, TiO2, P2O5, LREE and HFSE between the J3-T and J3-H samples in Group 2 suggest that the two subgroups were not comagmatic in origin. Intensive ferromagnesian fractionation (e.g. olivine and clinopyroxene) is essential to explain their low MgO, FeOT, CaO and other compatible elements

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Fig. 8. 1000/Sr v. 87Sr/86Sr(i) (a, b) and 100/Nd v. 1Nd(t) (c, d) for the Mesozoic Yanshan belt mafic rocks, showing the possible effect of crustal contamination and AFC during magma evolution. (a) and (c) show bulk contamination of possible crustal components; (b) and (d) show AFC trends. The calculation endmembers are: the NCB UCC: Sr 210 ppm, Nd 40 ppm, 87Sr/86Sr(i) ¼ 0.716; 1Nd(t) ¼218 (our unpublished data from Cenozoic sedimentary rocks in the NCB); the supposed NCB LCC–MCC: Sr 350 ppm, Nd 20 ppm, 87Sr/86Sr(i) ¼ 0.708; 1Nd(t) ¼ 230 (Jahn & Zhang 1984; Zhou et al. 2002; Liu et al. 2004); 98LP-2: Sr 1323 ppm, 87Sr/86Sr(i) ¼ 0.7052; 20XSH-41: Nd 52 ppm, 1Nd(t) ¼ 27.6. In the AFC modelling, the assimilation rate (r) is assumed to be 0.25. (See details in the text.) Tick marks on the AFC curves represent increments of 0.1 remaining magma fraction (F).

(e.g. Cr and Ni). Plagioclase fractionation is insignificant, as both subgroups have fairly high Sr contents. For the K1-D samples in Group 3, a fractional assemblage of olivine þ clinopyroxene + ilmenite + apatite is required to account for the rapid decrease in MgO, FeOT, CaO, TiO2 and P2O5 following magma evolution. Plagioclase fractionation is subordinate, as both Al2O3 and Sr show little variation over the wide range of SiO2. The K1-H rocks, which have the lowest MgO, and highest P2O5 and Th, require intensive ferromagnesian and plagioclase fractionation and subordinate apatite fractionation before extrusion to surface.

Source characteristics In recent years, geochemical investigations of the Mesozoic mafic rocks, and of Cenozoic

basalt-hosting granulite and pyroxenite xenoliths, have revealed the existence of highly enriched but heterogeneous mantle sources beneath the northern NCB (e.g. Zhang et al. 1998; Yan et al. 2000; Shao et al. 2001; Zhou et al. 2001, 2002; Xu 2002; Zhang et al. 2003; Li et al. 2004; Liu et al. 2004). Debate still continues about the timing and mechanism of the enrichment event(s). Possible models are: Mesozoic lithospheric delamination to account for lithospheric thinning and incompatible element enrichment in the NCB (e.g. Gao et al. 1998, 2004; Liu et al. 2004); slab-modified lithospheric mantle to explain the petrogenesis of high-Mg andesites (Zhang et al. 2003); multiple-stage enrichment processes related to recycling of NCB crust to form the pyroxenite xenoliths (Xu 2002); and lower crust –lithospheric mantle interaction to explain the radiogenic Sr and unradiogenic Nd and Pb isotopic compositions in the granulite xenoliths of

GEOCHEMISTRY OF MESOZOIC MAFIC VOLCANIC ROCKS

Mesozoic ages (Zhang et al. 1998; Zhou et al. 2002; Wilde et al. 2003). However, available Sr– Nd–Pb isotope data for mafic rocks with ages from 245 to 110 Ma, including basalts, high-Mg andesites, mafic cumulates and granulite xenoliths hosted in Cenozoic basalts, indicate radiogenic Sr and unradiogenic Nd and Pb isotopic compositions (Zhang et al. 1998, 2003; Li et al. 2001, 2004; Shao et al. 2001; Zhou et al. 2001, 2002; Gao et al. 2004). These data at least indicate that before the early Mesozoic an enrichment event occurred during the evolution of lithospheric mantle beneath the Yanshan belt. Because of the insignificant role of crustal contamination or AFC during magma evolution, the strong depletion in Nb–Ta and Th–U, and moderately radiogenic Sr and unradiogenic Nd isotopic compositions in the Yanshan belt, the mafic rocks require their derivation from a LILE- and LREE-enriched lithospheric mantle. The adakitic trace element features in the J3-T basaltic andesites suggest a significant contribution of slab melts in the melting mantle source (Rogers et al. 1985; Saunders et al. 1987; Calmus et al. 2003; Martin et al. 2005), rather than interaction between melts derived from delaminated eclogitic crust and upper mantle as suggested by Gao et al. (2004). We thus relate the enrichment process to the subduction of the palaeo-Asian Ocean during late Palaeozoic– early Mesozoic time (e.g. Sengo¨r & Natal’in 1996; Robinson et al. 1999; Shi et al. 2003; Miao et al. 2004). It is noteworthy that the considerable U –Th troughs in the spidergrams and relatively higher TiO2 contents in the least evolved samples (e.g.

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20XSH-41, TiO2 2.11%) from the Yanshan belt are distinguishable from the results for most arc basalts (e.g. Arculus 1994; Hunter & Blake 1995; Macdonald et al. 2000; Jolly et al. 2001; Plank 2005). Based on Th/La values of global arc basalts and the subducting sediments, Plank (2005) attributed the Th/La ratios in arc basalts to mixing between the underlying mantle wedge and the subducting sediments. He further pointed that where the subducting sediment has Th/La . 0.2, the arc basalts generally have Th/La . 0.15, and where the sediment has a value ,0.2, the arc basalts commonly have a value ,0.15 (Plank 2005). The Mesozoic Yanshan belt mafic volcanic rocks have low Th/La (0.039–0.10), with some samples from the J3-T and K1-D having even lower Th/La than N-MORB (Th/La ¼ 0.048, Sun & McDonough, 1989), calling for involvement of extremely low-Th/La crustal materials in their melting source, such as lower crustal granulite xenoliths or CaO- and P2O5-rich marine sediments (Plank 2005, and references therein; see also Huang & Frey 2005). The other samples, which have higher Th/La ratios than N-MORB, require the addition of relatively high-Th/La crustal rocks (Fig. 9), such as terrigenous sediments and/or continental upper crust. However, the generally low Th/La ratios in the Mesozoic Yanshan belt mafic rocks suggest that the involved crustal components also have a bulk low Th/La (e.g. ,0.2) according to the Th/La estimation from modern arcs (Plank 2005). The most likely candidates are the Archaean high-grade metamorphic terranes or lower/middle crust of the NCB, which have low Th/La and radiogenic Sr and unradiogenic Nd isotopic

Fig. 9. Th/La v. 1Nd(t) (a) and La/Nb (b) for the Mesozoic Yanshan belt mafic rocks. The extremely low Th/La (,0.048, N-MORB) in some samples suggests the contribution of a low-Th/La, low 1Nd(t) and high La/Nb component (Sun & McDonough 1989), whereas the other samples require a high-Th/La and La/Nb, and low-1Nd(t) endmember in the melting source. (See discussion in the text).

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compositions (Jahn & Zhang 1984; Huang et al. 2004; Liu et al. 2004; Zhou et al. 2002). One possibility to introduce the low-Th/La crustal materials into the melting mantle source is that they were eroded and deposited as terrigenous sediments and were dragged down into the mantle wedge together with the subducting palaeo-Asian slab during the late Palaeozoic to early Mesozoic (e.g. Sengo¨r & Natal’in 1996; Robinson et al. 1999; Shi et al. 2003; Miao et al. 2004). Another possibility is that the palaeo-Asian Ocean was an intracontinental oceanic basin, in which the NCB lower continental crust remained. For instance, there exist inherited Archaean zircons in the mafic components of the Hegenshan ophiolite (e.g. Miao et al. 2006). During the slab subduction, both the oceanic crust and the NCB lower crustal materials or recycled Archaean high-grade metamorphic terranes were dragged down into the mantle. Interaction of these ancient crustal materials with the mantle beneath the northern NCB would form a mantle reservoir enriched in LREE and LILE but of low Th and U, radiogenic Sr, and unradiogenic Nd and Pb isotopic compositions.

Comparison with the NCB lithospheric mantle The geochemical differences between the late Mesozoic NCB interior and the Yanshan belt mafic lavas (Figs 4, 6 and 7) also indicate mantle heterogeneity between cratonic and circumcratonic regions. Geochemistry of the late Mesozoic NCB interior mafic rocks reveals the existence of a chemically refractory and long-term LILE- and LREE-enriched mantle reservoir, which is a feature of subcratonic lithospheric mantle (Guo et al. 2001, 2003; Zhang et al. 2004). In the following discussion, we compile available data to show the differences between the SCLM beneath the two regions for major element considerations. An approach commonly used for comparison of basalt data to distinguish mantle domains is to extrapolate the less evolved rocks to 8% MgO, beyond which fractionation is believed to be largely dominated by olivine (Klein & Langmuir 1987). As the Mesozoic Yanshan belt mafic rocks are fairly well evolved in terms of major element compositions (e.g. Mg-number ,60, and mainly 45–57 for samples with MgO . 4.5%), the assumption that the fractionating assemblage remains constant when extrapolated back to 8% MgO is difficult to verify (e.g. Garland et al. 1996). Where obvious inflections in the trends occur as a result of changes in the fractionating assemblage a line is fitted to the most primitive portion of the trend, and obvious outliers are

eliminated. Extrapolation of the Mesozoic Yanshan belt basalts and the late Mesozoic NCB interior mafic rocks to 8% MgO has been carried out by fitting a simple correlation line; only samples with a MgO range of 4.5 –10% are regressed, following the method and rationale of Turner & Hawkesworth (1995). The calculation results are listed in Table 5. The Yanshan belt basalts and mafic dykes (Shao et al. 2001) span a Si8 range of 43.9 –51.3%, Fe8 range of 9.3–13.8% and Ti8 range of 1.8–3.3%. Compared with the late Mesozoic NCB interior counterparts (Guo et al. 2001, 2003, and unpublished data), they have higher Fe8, Ti8 and lower Si8 (Fig. 10a and b). In contrast to the late Mesozoic NCB mafic rocks, which have moderately varied Na8 (2.2–3.3%), the Yanshan belt rocks have a relatively larger Na8 range from 1.6 to 4.0% (Fig. 10c), a result of either surface alteration or the incorporation of residual Na in the jadeite component of clinopyroxene during melting (e.g. Klein & Langmuir 1987; Turner & Hawkesworth 1995; Garland et al. 1996). As shown in Figure 10b and d, there are no coherent variations between the Mesozoic mafic rocks from the two areas; in particular, there is a Ti8 gap from 0.9 to 1.8%. Experimental results from peridotite melting reveal that the Fe content of a melt is positively correlated with pressure and source composition; thus a melt either at high pressures or derived from a fertile source will have high Fe (e.g. Hirose & Kushiro 1993; Baker & Stolper 1994). In contrast, the Si content is negatively correlated with melting pressure and increases in the presence of fluid. The Ti content is a gauge of source fertility and is sensitive to melt removal. Na mainly reflects the degree of melting, as it acts as an incompatible element under mantle conditions. At high pressures, Na is less incompatible, as it is retained in the residue incorporated in the jadeite component of clinopyroxene. Compared with experimental results (e.g. Kushiro 1990; Hirose & Kushiro 1993; Baker & Stolper 1994), the higher Si8, Fe8 and Ti8 in the Yanshan belt basalts than those of the late Mesozoic NCB interior mafic rocks can be attributed to either variation in melting conditions (e.g. pressure and melting degree) or source heterogeneity, or both (Fig. 10). Because Th and U are highly incompatible elements during melting, Th and U contents in basalts are sensitive to the extent of melting. As can be seen in the Th v. U diagram (Fig. 11), both the Yanshan belt and the NCB interior mafic rocks with a MgO range of 4.5–10% have comparable Th and U contents, suggesting that melting degree was not a major factor to account for the Ti8 gap between the mafic rocks from the two

19 15

Number of samples

R2 0.51 0.04

Si8 average + 1s 52.38 + 0.85 47.50 + 5.24

8.17 + 0.70 10.64 + 0.41

Fe8 average + 1s

Correction of the data was carried out using the equation X8 ¼ (slope  8 þ Xsample) 2 (slope  MgOsample).

The NCB interior The Yanshan belt

Locality 0.03 0.11

R2

0.81 + 0.01 2.51 + 0.16

Ti8 average + 1s

0.45 0.22

R2

Table 5. Averages at wt% MgO for selected major elements of the late Mesozoic mafic rocks from the NCB interior and the Yanshan belt

2.56 + 0.10 3.24 + 0.97

Na8 average + 1s

0.77 0.05

R2

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Fig. 10. Si8 v. Fe8, Ti8 and Na8 and Na8 v. Ti8 for the Mesozoic Yanshan belt basalts and late Mesozoic NCB mafic rocks, showing the possible effect of melting conditions and source composition. The higher Fe8 and Ti8 but lower Si8 contents in the Mesozoic Yanshan belt basalts reflect a relatively fertile source in the origin. Samples with a MgO range of 4.5– 10% are selected for 8% MgO extrapolation. Data sources are as in Figure 3.

Fig. 11. Th v. U for Mesozoic Yanshan belt basalts, showing possible effect of magmatic processes on magma evolution. Samples with a MgO range of 4.5– 10% are selected for comparison. Data source are as in Figure 3.

regions, even if these melts had been derived from a common source. We thus consider the Ti8 gap as a result of source difference; that is, there was a more fertile mantle source beneath the Yanshan belt than beneath the NCB interior. Irrespective of the effect caused by melting pressure on melt composition, the relatively fertile mantle source would also explain the lower Si8 and higher Fe8 in the Yanshan belt basalts. In summary, the lower Si8 but higher Fe8 and Ti8 in the mafic rocks from the Yanshan belt compared with those in the NCB interior indicate that they were probably derived from a chemically fertile mantle reservoir, different from the Mesozoic chemically refractory mantle reservoir beneath the NCB interior (Guo et al. 2001, 2003). Such a difference is also reflected by the generally higher Sr isotope ratios in the Yanshan belt mafic lavas, which suggest modification of their source by high-Rb/Sr and high-87Sr/86Sr materials (e.g. ancient continental crust or

GEOCHEMISTRY OF MESOZOIC MAFIC VOLCANIC ROCKS

terrigenous sediments). Accordingly, the different mantle domain beneath the Yanshan belt compared with the NCB interior is in accord with lithospheric mantle evolution in circumcratonic terranes, where subduction-related crustal recycling plays an important role in the refertilization of SCLM (e.g. Boyd 1989; Menzies 1989).

Geodynamic implications for Mesozoic tectonomagmatism According to the above discussion, the Mesozoic Yanshan belt mafic lavas were derived from a heterogeneous SCLM that had been previously enriched in LILE and LREE. Although the Mesozoic mafic volcanism is subordinate when compared with the predominant intermediate –felsic magmatism, generation of these mafic magmas provides important constraints on the dynamics of melting within the SCLM and geodynamic models for lithospheric evolution through Mesozoic time. Both numerical modelling and thermal considerations suggest that in the absence of plume activity, basaltic generation in continental settings is linked with either lithospheric extension or slab subduction (e.g. McKenzie & Bickle 1988; Arculus 1994; Kerr 1994; Hawkesworth et al. 1995; Hunter & Blake 1995; Macdonald et al. 2000; Fan et al. 2003). The predominant intermediate –felsic lavas and subordinate mafic rocks in the Yanshan belt are compositionally different from plumerelated large igneous provinces, which mainly comprise basalts, thus precluding the role of plume impact. On one hand, the 240– 250 Ma peak metamorphism from the Lesser Hinggan Mountain to the Xilin Gol metamorphic belt and the emplacement of Indo-Sinian alkaline complexes along the northern NCB marked the end of orogeny related to the subduction of the palaeo-Asian Ocean (e.g. Yan et al. 2000; Shi et al. 2003; Miao et al. 2004). An origin related to the palaeo-Asian Ocean subduction seems unsuitable for these mafic rocks. On the other hand, the occurrence of a series of accretionary prisms from Jurassic to Cenozoic time in the Japan arcs suggests the role of multistage oceanic slab subduction during the growth of the East Asian continent (e.g. Faure & Natal’in 1992; Taira 2001), such accretionary complexes are absent in the Yanshan belt and its surrounding regions. Palaeomagnetic surveys of the Pacific Ocean and the Eurasia continent have indicated that the westward subduction of the palaeo-Pacific Ocean began at c. 100 Ma (Engebretson et al. 1985; Northrup et al. 1995). The Mesozoic (180 –110 Ma) Yanshan belt mafic volcanism predates the westward Pacific subduction event. Similarly, an origin related to the palaeo-Pacific Ocean subduction is also inapplicable

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to interpret the Mesozoic Yanshan belt mafic rocks. Accordingly, a multistage lithospheric extension model is envisaged to explain the basaltic generation in the Yanshan belt through Mesozoic time. The Group 1 mafic volcanic rocks. The Group 1 mafic rocks in the Yanshan belt are sporadically distributed along the Xinglong–Kuancheng and Chicheng –Luanping faults, and crop out in small volumes (HBGMR 1989). Geochemical and geothermal considerations suggest that such small volumes of SCLM-derived mafic melts could be produced by either thermal perturbation or local lithospheric extension. Geometric surveys and thermal chronology of structural deformation identified a pre-180 Ma south-vergent thrust event prior to the eruption of the Nandaling Fm. basalts (Fig. 12a). Despite the controversy about the dynamics of this event (e.g. Wang 1996; Deng et al. 1999; Davis et al. 2001), the nearly east – west-trending lineation on the thrust faults suggests that their formation was related to a north– southdirected compressional stress field, either as a consequence of the collisional suturing of Palaeozoic Mongolian arcs against an Andean-style continental arc along the northern NCB (Shi et al. 2003; Miao et al. 2004), or as an expression of a back-arc, foreland fold and thrust belt of US Cordilleran type formed during southward subduction of the Mongolia- Okhotsk Ocean beneath the NCB (Fig. 12a; Davis et al. 2001). Subsequent extension in response to either post-collisional stress relaxation or lithospheric readjustment triggered asthenospheric upwelling and partial melting of the previously enriched sources to generate these mafic magmas (Fig. 12b).

The Group 2 basaltic lavas In essence, the Group 2 volcanism in the northern NCB is characterized by extrusion of predominant intermediate –felsic lavas and minor basaltic rocks, such as the Tiaojishan Fm., which is mainly composed of calc-alkaline andesites and dacites, with subordinate mafic rocks (e.g. Bao et al. 1995; Li et al. 2001). Extensive melting of crustal rocks requires a high heat supply as a result of either crustal thickening or basaltic underplating. Unlike syntectonic peraluminous intermediate – felsic rocks, which can be produced through melting of middle –upper crustal rocks by means of crustal thickening for accumulation of high heat productivity radioactive elements (e.g. Thompson 1982; Ellis & Thompson 1986), the Group 2 intermediate –felsic lavas in this region mainly have metaluminous to calc-alkaline affinities, and were derived from the lower crust and/or

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Fig. 12. A schematic illustration of the Mesozoic tectonic evolution of the Yanshan belt. (a) At c. 240 Ma, the northern NCB was involved in the collision between the Mongolian Block and NCB (Shi et al. 2003; Miao et al. 2004); resulting in a pre-180 Ma contractional deformation event (Zhao et al. 1990; Davis et al. 2001). (b) At c. 180 Ma, local extension in response to post-compressional stress relaxation or lithospheric readjustment along the pre-existing major faults (e.g. the GXPF), resulting in the eruption of sporadic Nandaling Fm. basalts (e.g. Bao et al. 1995; Li et al. 2004). (c) In the Middle Jurassic, collision between the North China– Mongolian Block and the Siberian plate (e.g. Zhao et al. 1990) resulted in the north– south directed contractional deformation in the northern NCB. (d) In the Late Jurassic, post-collisional extension and large-scale volcanic eruption took place in the NE China Fold belt (Fan et al. 2003) and the northern NCB (Group 2), and gave rise to syntectonic kilometre-scale sediment deposition (e.g. Davis et al. 1998, 2001; He et al. 1998; Zheng et al. 1999; Cui et al. 2002). (e) In the Early Cretaceous, the Group 3 volcanic rocks were extruded under an extensional regime in response to lateral escape tectonics related to surrounding plate interactions.

underplated mafic cumulates (e.g. Li et al. 2001). This required significant lithospheric extension to trigger basaltic underplating to reheat the lower – middle crust (Bergantz 1989), and the scarce exposure of basaltic lavas might be a result of mafic cumulates emplaced at crustal levels, as

indicated by magmatic granulite-facies rocks at Hannuoba (Fan et al. 1998; Zhang et al. 1998; Zhou et al. 2002; Wilde et al. 2003). Because the collision between the North China–Mongolian Block and the Siberian plate was terminated by the end of the middle Jurassic (e.g. Zhao et al.

GEOCHEMISTRY OF MESOZOIC MAFIC VOLCANIC ROCKS

1990), contractional deformation took place along the Yanshan – Yinshan belt (e.g. Zhang & Song 1997; Zheng et al. 1999; Davis et al. 2001), for example, giving rise to the Chengde –Xinglong thrust and Shihetang nappes (Fig. 12c). Thermal decay and decrease in the amount of lithospheric extension triggered small-scale melting of the mantle source to produce shoshonitic magmas in the Houcheng Fm. Both the structural deformation styles and the metaluminous affinities of magmatic activity in this region are essentially similar to those in a post-collisional setting (e.g. Turner et al. 1996; Miller et al. 1999; Fan et al. 2003). In combination with the almost east –west-trending distribution of the volcanic sequences and structural alignment of the sedimentary basins, the lithospheric extension also had a north –south-trend (Fig. 12d), either as a consequence of thermal perturbation caused by the far-field effect from the collision of the two continents or as an expression of post-collisional collapse.

The Group 3 basalts By the end of the Jurassic or at the beginning of the Cretaceous, there existed a regional angular unconformity between the Houcheng or Tuchengzi Fm. sediments and the Zhangjiakou or Donglingtai Fm. felsic lavas. The cause of this tectonic event is still unclear; it may have been a result of gravitational collapse of the Yanshan belt or a consequence of strike-slip extension induced by the northward motion of the palaeo-Pacific Ocean. The subsequent lithospheric extension reached its maximum, accompanied by voluminous eruption of magmas (the Zhangjiakou and Huajiying Fms. in the Jibei region, the Yixian Fm. in the Liaoxi region, and the Donglingtai and Donglanggou Fms. in the Xishan area and within the NCB (Fig. 12e; e.g. Bao et al. 1995; Chen et al. 1997; Zhou et al. 2001; Cui et al. 2002; Guo et al. 2003). During this period, the volcanism was also strongest throughout eastern China, forming a NE– SW-trending magmatic belt thousands of kilometres in extent. In recent years, structural studies of late Mesozoic deformation in eastern China have indicated a prevalent extensional regime associated with strike-slip displacement and lateral extrusion in response to surrounding plate interactions (Xu et al. 1987; Yin & Nie 1993; Hacker et al. 2000; Ratschbacher et al. 2000; Wang et al. 2003). For instance, the NCB was subjected to southward and northeastward compression from the Siberian Plate and the Lhasa –Qiangtang Block, respectively (e.g. Menzies et al. 1993; Yin & Nie 1993; Hacker et al. 2000; Ratschbacher et al. 2000; Guo et al. 2003; Wang et al. 2003). These two indentors led to NE-directed lateral escape of the lithosphere.

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Simultaneously, the rapid northward movement of the palaeo-Pacific Ocean and the induced largescale strike-slip displacement of lithosphere-scale faults (Engebretson et al. 1985; Northrup et al. 1995), such as the Tan-Lu wrench fault system (Xu et al. 1987), provided a focus for strike-slip deformation along the East Asian continental margin. The resultant force was favourable for rapid dispersion of extrusive materials, the development of extensional tectonics, and extensive magmatism in the Yanshan belt and its surrounding areas.

Conclusions The Mesozoic Yanshan belt mafic lavas, characterized by significant LILE and LREE enrichment but Nb –Ta and Th –U depletion, and moderately radiogenic Sr and unradiogenic Nd isotopic compositions, were derived from a heterogeneous lithospheric mantle that had been previously enriched in LILE and LREE, probably as a consequence of recycling of the ancient NCB lower crustal materials related to the palaeo-Asian Ocean subduction. The high Ti8 and Fe8, and low Si8 relative to the late Mesozoic NCB interior mafic rocks indicate that the Yanshan belt mafic lavas were probably derived from a relatively fertile source, different from the chemically refractory, lithospheric mantle within the NCB. Such a difference is in accord with mantle evolution beneath circumcratonic terranes, where recycled crustal materials play an important role in refertilization of mantle lithosphere. Combining the Mesozoic magma petrogenesis with the crust – lithosphere deformation history, the multistage intracontinental tectonomagmatic episodes occurring in the Yanshan belt were responses to surrounding plate interactions with the lithospheric evolution of the NCB continental margins. The authors would like to thank L. Qi for his help with ICP-MS, and Y. Liu and X. R. Liang for Sr– Nd isotope analyses. H. F. Zhang and Y. J. Wang are thanked for discussion. Constructive reviews and helpful suggestions from A. Saunders, M. Lamb and T. Vislova help to greatly improve the manuscript. This study was financially supported by Chinese Academy of Sciences (KZCX1-107) and the National Natural Science Foundation (40034343).

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Origin of the Mesozoic magmatism in the North China Craton: constraints from petrological and geochemical data B. CHEN1, M.-G. ZHAI2 & W. TIAN1 1

School of Earth and Space Sciences, Peking University, Beijing 100871, China (e-mail: [email protected])

2

Institute of Geology and Geophysics, Chinese Academy of Sciences, Beijing 100029, China Abstract: Voluminous plutonic and volcanic rocks were emplaced in the eastern part of the North China Craton (NCC) in the Mesozoic. The Mesozoic igneous rocks include a variety of rock types ranging from monzogabbroic, through monzonitic to monzogranitic, and locally to syenitic. Monzonitic rocks are dominant, and frequently contain mafic enclaves (dioritic in composition). The principal geochemical signatures of these Mesozoic rocks include high-K calc-alkaline to shoshonitic affinity, high Sr –Ba abundances and high Sr/Y, La/Yb, and highly enriched Sr– Nd isotopic compositions with 1Nd(t) ranging from 28 to 220 and ISr from 0.7053 to 0.710. Zircon sensitive high-resolution ion microprobe dating reveals that these Mesozoic rocks formed between 180 Ma and 120 Ma, but are predominantly confined to a narrow range of 135–127 Ma. The sudden surge of Mesozoic magmatism was genetically linked to the upwelling of asthenosphere in a back-arc extensional regime that was caused by subduction of the palaeo-pacific plate beneath the eastern NCC. Upwelling of hot asthenospheric mantle material triggered partial melting of enriched subcontinental lithosperic mantle, generating voluminous mafic magmas. The mafic magmas underplated in the lower crust and sparked melting of the latter, producting granitic melts. We suggest that the Mesozoic rocks in the NCC probably originated from mixing between the coeval mafic and granitic melts, followed by fractionation of ferromagnesian phases and subordinate plagioclase, rather than from melting of mafic lower crust as previously suggested by many others.

The North China Craton (NCC) is one of the oldest continental nuclei in the world (Jahn et al. 1987), with basement rocks of Archaean to Palaeo-Proterozoic gneisses (tonalite–trondhjemite– granodiorite; TTG). It was stabilized in late Palaeo-Proterozoic time at c. 1.85 Ga, and much of it remained stable to Trassic time. However, the eastern part of the NCC has been tectonically active since the early Mesozoic, and experienced intense tectonothermal events as shown by the development of large-scale sedimentary basins and widespread magmatism (and related metallogeny). Many published data (e.g. Menzies et al. 1993; Menzies & Xu 1998; Xu 2001) indicate that the chemical and isotopic signature of the subcontinental lithospheric mantle (SCLM) beneath the eastern NCC changed isgnificanlty after the Mesozoic tectonothermal events. The occurrence of diamonds in Palaeozoic kimberlites implies the presence of thick (150 –220 km), old and refractory lithospheric keel underlying the craton in preMesozoic time, whereas mantle xenoliths in Cenozoic basalts suggest a thin (60–120 km), young and fertile oceanic-type lithosphere in the Cenozoic. Such a change of lithospheric nature implies a significant erosin of the old SCLM during the

Mesozoic (Menzies & Xu 1998; Xu 2001; Gao et al. 2002), and distinguishes the NCC from most other old cratons such as South Africa (Boyd & Gurney 1986) and Western Australia (Anderson et al. 1992), which are characterized by the existence of thick Archaean mantly lithosphere keels. What is the origin and cause of the sudden surge of Mesozoic magmatism in the NCC, and in what kind of geodynamic setting did it occur? What is the relationship between the Mesozoic magmatism and lithosphere evolution or thinning? These issues have been subjects of extensive research of both domestic and international workers in recent years (e.g. Menzies & Zu 1998; Xu 2001; Zheng et al. 2001; Gao et al. 2002; Zheng et al. 2002; Chen & Zhai 2003; Chen et al. 2003a, b). The lithospheric processes in eastern China have had profound effects on the tectonics and Mesozoic magmatism of this region. In this paper, we compile a large elemental and isotopic dataset for the igneous rocks from across the eastern NCC, which, in combination with our new petrological data, is used to understand the sources of these rocks, and the relationship between the Mesozoic magmatism and lithosphere thinning.

From: ZHAI , M.-G., WINDLEY , B. F., KUSKY , T. M. & MENG , Q. R. (eds) Mesozoic Sub-Continental Lithospheric Thinning Under Eastern Asia. Geological Society, London, Special Publications, 280, 131– 151. DOI: 10.1144/SP280.6 0305-8719/07/$15 # The Geological Society of London 2007.

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General geology and ages of Mesozoic magmatism As seen in Figure 1, the NCC is located in the eastern part of the Euro-Asian continent, bounded to the north by the southern margin of the Central Asian Orogenic Belt (CAOB) and to the south by the Dabie –Sulu ultrahigh-pressure (UHP) belt, which formed during the process of a Triassic collision between the Yangtze Block and the NCC (Li et al. 1993; Jahn et al. 1999). The Sulu UHP belt was originally the eastern extension of the Dabie UHP belt, but has been truncated and moved to the north along the Tanlu fault (Fig. 1). The basement rocks of the NCC are Archaean to Palaeo-Proterozoic gneisses of amphibolite to granulite facies (Jahn et al. 1987). They are unconformably overlain by thick sequence of Mid- to Neo-Proterozoic quartzites and limestones, suggesting that the craton was stabilized by the Palaeo-Proterozoic. Based on geophysical data, Ma (1987) identified a giant south –north gravity lineament (NSGL) that divides the NCC into two different tectonic domains at present (Fig. 1). The region to the west of the gravity lineament is characterized by large negative Bouguer gravity anomalies and a thick (150 –220 km) lithosphere, in contrast to the weakly negative to positive Bouguer gravity anomalies, high heat flow and a relatively thin lithosphere (60 –120 km) in the region to the east. Mesozoic magmatism, together with the development of large-scale sedimentary basins, occurs mainly in the region east of the gravity lineament, in contrast to the minor magmatism and basin development in the region west of the gravity lineament (Fig. 1). Mesozoic

Fig. 1. Geological sketch map of East China showing distribution of the Mesozoic magmatism (grey fields). NSGL, north–south gravity lineament (Ma 1987); CAOB, Central Asian Orogenic Belt. (See text for further details.)

magmatism is characterized by intrusion of voluminous intermediate to felsic rocks (dominantly monzonitic) and related mafic rocks, and eruption of widespread volcanic rocks. As seen in Figure 1, Mesozoic igneous rocks of the NCC occur mainly in three places: the Taihang–Yanshan orogen, the Jiaodong Peninsula, and the Dabie UHP belt. In this paper, we compile a large geochemical and isotopic dataset for igneous rocks mainly from these three representative regions. The Taihang–Yanshan orogen lies in the western part of East China, stretching along the NSGL (Fig. 1). Mesozoic intrusions of the orogen are dominantly monzonitic to quartz monzonitic in composition, and are accompanied by a few related mafic (monozogabbroic to monzodioritic) and syenitic plutons (Chen et al. 2003a, 2004). Chen et al. (2005) reported sensitive highresolution ion microprobe (SHRIMP) zircon U –Pb ages of 132 + 2 Ma and 129 + 2.6 Ma for a quartz monozonite pluton and a monzonite pluton in the orogen, respectively. A slightly older age of 138+2 Ma was also reported by Chen et al. (2005) for a mafic pluton in the orogen. These ages are consistent with the age data (130 – 140 Ma, zircon U – Pb) reported for these monzonitic rocks by Davis et al. (1998). The Jiaodong Peninsula is in the easternmost part of the MCC (Fig. 1). Igneous rock types of Jiaodong are similar to those of the Taihang– Yanshan orogen, essentially being monzonitic in composition, accompanied by subordinate granitic rocks and minor mafic plutons. Many high-quality age data for rocks from Jiaodong were published recently. Most plutons were dated in the range 145–125 Ma using SHRIMP zircon U –Pb methods (Yang et al. 2004), although a few plutons with older ages (160 –180 Ma; Wang et al. 1998) were also found in the area. Voluminous Mesozoic magmas were emplaced in the Dabie UHP belt, which lies in the southern part of NCC (Fig. 1). The main rock types include monzonite, quartz monzonite and monzogranitic rocks, which also are accompanied by minor, coeval mafic rocks as in the Taihang– Yanshan orogen and the Jiaodong Peninsula. Hacker et al. (1998) reported a large dataset of zircon U –Pb ages (SHRIMP method) for these rocks, which cluster in the range 137–126 Ma, similar to that for the Taihang–Yanshan and Jiaodong rocks. Therefore, the majority of the Mesozoic magmatism in the NCC took place in the period 127–138 Ma.

Mafic enclaves and plagioclase compositions One of the striking features of the NCC Mesozoic magmatism is the occurrence of minor, but

ORIGIN OF THE MESOZOIC MAGMATISM

widespread mafic enclaves that are hosted in the intermediate to felsic plutons. Mafic enclaves are typically centimetres to tens of centimetres in diameter. The lobate to cuspate contacts against host rocks (Fig. 2a) suggest that the two rock units coexisted as contemporaneous magmas, and the absence of chilled margins in enclaves suggests a small temperature contrast between the coexisting

Fig. 2. (a) Mafic enclaves in monzonitic pluton, showing lobate to cuspate contacts with host rocks, (b) Local concentration of mafic enclaves. (c) Kfs phenocryst in a mafic enclave; it has chemical composition similar to that of the host rocks. (Width of field of view is 8 cm.)

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magmas. Interaction (or hybridization) between mafic enclave and surrounding felsic magma is apparent from gradational contacts, suggesting processes of magma mixing or mingling. Locally, mafic enclaves are highly concentrated (c. 30% of the outcrop; Fig. 2b) and in such cases the matrix typically shows irregular leucocratic patches and areas rich in fine-grained minerals. This feature suggests dissemination of enclave material, which yields heterogeneous hybrid rocks containing wispy schlieren and clots of fine-grained mafic material that are heterogeneously distributed throughout the host rocks. Typically, mafic enclaves are monzodioritic in composition, being made up of plagioclase (40–50%), K-feldspar (10–15%), hornblende (30%), biotite (5–10%), pyroxene (5–10%), +quartz, and accessory titanite, magnetite, and apatite. Their texture varies from equigranular, fine-grained to porphyritic, but is exclusively igneous. Alkali feldspar phenocrysts, which are texturally and compositionally identical to those in the host rocks, are observed in some enclaves (Fig. 2c), suggesting transfer of phenocrysts before magma was solidified and the coevality of mafic enclaves and host rocks. Mafic enclaves are interpreted to represent mingling or mixing of basaltic magma into the intermediate to felsic magma during influx before the latter was solidified. Magma mixing is also indicated by the compositional and textural disequilibrium of plagioclase crystals in both mafic enclaves and host rocks. Figure 3a and b shows a typical resorbed calcic core (An58 – 65) surrounded by an abrupt sodic rim (An35 – 38; see Table 1 for the microprobe analysis of plagioclase cores and rims). This differs from ‘normal’ compositional zoning of plagioclase, which has a gradational, rather than abrupt, change of chemical composition. We note that plagioclase cores typically are eroded and show embayments, but otherwise are euhedral (Fig. 3a and b), showing relatively homogeneous, calcic composition, with An ranging from 65 to 58. The mantles of plagioclase, however, are clean, showing apparently lower An number (An35 – 38) than the cores. We interpret the calcic cores to have formed by two possible processes, as follows. (1) A plagioclase crystal that crystallized from felsic magma was entrapped in basic magma and subsequently experienced partial melting (and heterogeneous diffusion) as a result of the high temperature of the magma, leaving a more calcic restite. (2) The eroded, calcic cores were originally crystallized from mafic magma, and then were entrapped in hybrid magma formed through magma mixing between mafic and granitic melts. In either case, the relatively sodic mantle represents magmatic overgrowth from hybrid magma (Blundy & Sparks 1992; Janousek et al. 2004; Kemp 2004). The preservation of the compositional and textural

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Fig. 3. Compositional and textural disequilibrium of a plagioclase crystal that has a resorbed calcic core (An58 – 65; dark grey) surrounded by an abrupt sodic rim (An35 – 38; light grey). Microprobe analyses of plagioclase are shown in Table 1. (Width of field view is 2.5 cm.)

disequilibrium is due to rapid crystallization and incomplete mixing. The disequilibrium shown by plagioclase xenocrysts, together with the presence of widespread mafic enclaves within plutons, strongly suggests a process of magma mixing or mingling, which could have played a dominant role in generating the Mesozoic magmas in the NCC. This model is compatible with the elemental and isotopic data for the Mesozoic rocks as well as mafic enclaves, as described below.

Major and trace element data Tables 2, 3 and 4 show representative chemical data for the igneous rocks from Taihang–Yanshan, Jiaodong and Dabie, respectively. Also shown in the tables are chemical data for typical mafic enclaves. Figure 4 is a classification diagram in which a large dataset for the Mesozoic rocks from the NCC were plotted. Rocks from Taihang– Yanshan, Dabie and Jiaodong are indistinguishable

Table 1. Microprobe analysis of plagioclase Spot:

1

2

3

4

5

6

7

8

9

SiO2 Al2O3 TiO2 FeO MgO CaO Na2O K2O Total

60.746 24.843 0.032 0.178 0.008 5.679 7.861 0.351 99.70

59.56 26.396 0.02 0.259 0.009 8.059 7.091 0.383 101.78

57.287 25.564 0.001 0.268 0.001 7.835 6.786 0.386 98.13

48.091 29.861 0.007 0.227 0.001 11.638 4.656 0.2 94.68

51.857 29.381 0.011 0.227 0.002 12.232 4.653 0.148 98.51

58.287 25.377 0.001 0.305 0.001 7.238 7.387 0.524 99.12

58.505 25.536 0.005 0.233 0.016 7.525 7.324 0.363 99.51

52.013 29.91

52.788 29.737 0.015 0.219

Cations (per 8 oxygens) Si 2.7052 2.6204 Al 1.3041 1.3687 Ti 0.0011 0.0007 Fe 0.0066 0.0095 Mg 0.0005 0.0006 Ca 0.271 0.3799 Na 0.6788 0.6049 K 0.0199 0.0215 Total 4.99 5.01 Ab 70.0 60.1 An 27.9 37.8 Or 2.1 2.1

2.6146 1.3753 0.0016 0.0005 0.3832 0.6005 0.0225 5.00 59.7 38.1 2.2

2.3136 1.6933 0.0003 0.0091 0.5999 0.4343 0.0122 5.06 41.5 57.3 1.2

2.3894 1.5956 0.0004 0.0087 0.0016 0.6039 0.4157 0.0087 5.02 40.4 58.7 0.8

2.6343 1.3519 0.0115

2.6316 1.3539 0.0002 0.0088

0.3505 0.6474 0.0302 5.03 63.0 34.1 2.9

0.3627 0.6387 0.0208 5.02 62.5 35.5 2.0

0.257 0.006 11.293 4.4 0.163 98.04 2.3966 1.6244 0.0099 0.0004 0.5576 0.3931 0.0096 4.99 40.9 58.1 1.0

12.655 3.628 0.193 99.24 2.4047 1.5967 0.0005 0.0083 0.6177 0.3205 0.0112 4.96 33.8 65.1 1.2

Plagioclase compositions were measured at Peking Univeristy, Beijing. Spot numbers 1– 5 are from Figure 3a; spot numbers 6– 9 are from Figure 3b. Data Source: Chen et al. 2006.

XG-4 gabbro

YM-2 gabbro

SiO2 47.39 46.61 TiO2 1.67 1.75 17.81 17.38 Al2O3 FeO 10.29 10. 08 MnO 0.17 0.17 MgO 5.30 6.20 CaO 7.20 8.37 Na2O 4.25 3.76 K2O 4.15 3.97 P2O5 1.28 1.06 LOI 0.3 0.50 Total 99.81 99.85 Mg-no. (ppm) 0.51 0.55 Rb 106 53.2 Sr 2878 1198 Ba 4039 1157 Zr 207 153 Hf 5.4 4.6 U 2.8 2.5 Th 12 4.9 Pb 8.7 9.1 Nb 71 12

Sample no: Rock type: 52.45 1.08 16.39 9.44 0.11 6.74 8.09 3.68 1.6 0.4 0.1 100.1 0.59 16.7 1488 1269 87.8 2.35 0.33 1.25 9.4 4.2

54.62 1.02 18.08 7.75 0.16 3.28 5.61 5.49 1.02 0.31 1.82 99.16 0.46 83 711 463 144 4.4 2.9 8.3 9.4 18

WA5 DH-7 monzo dio enclave

ZG-1 monz

56.49 59.02 0.79 0.95 18.16 16.57 7.75 6.55 0.11 0.09 2.82 3.6 6.28 5.41 4.63 4.22 2.59 3.16 0.39 0.34 0.4 0.67 100.41 100.58 0.42 0.52 60 74 1110 1293 847 1422 144 172 4.4 5.4 0.9 1.4 4.1 7.2 13 12 8.6 11

YM-5 enclave

WA-35 monz

YM-9 monz

63.24 59.16 64.09 0.81 0.56 0.65 16.17 18.47 15.57 5.25 5.45 5.2 0.1 0.08 0.08 2.46 1.81 2.42 3.71 4.47 4.02 4.67 4.3 3.88 3.53 4.59 3.74 0.25 0.34 0.2 0.42 0.43 0.49 100.61 99.66 100.34 0.48 0.40 0.48 82 80 102 745 1364 727 1432 2963 1145 198 173 150 6.7 4.8 5.1 1.4 1.6 2.6 9.1 7 17 10 11 14 16 14 13

DH-33 monz

Table 2. Chemical data for the Mesozoic rocks from the Taihang – Yanshan orogen

67.41 0.48 15.17 3.63 0.06 1.63 3.25 4.18 3.7 0.19 0.48 100.18 0.47 84 825 994 120 4.1 2.4 14 16 11

WA-16 qz. monzqz 67.29 0.49 16.2 3.24 0.06 1.1 2.53 4.69 3.7 0.2 1 100.5 0.4 54 832 1806 180 5.2 0.6 4.7 9.8 9.4

WA-7 monzqz

DH-9 monzqz

69.81 0.43 14.47 2.93 0.06 0.95 2.45 3.93 3.98 0.13 0.33 99.47 0.39 100 492 990 111 4.1 2.8 19 12 12

DH-22 monz

(Continued)

65.54 66.47 0.57 0.56 15.96 15.68 4.17 3.53 0.09 0.07 1.66 1.35 3.21 2.79 4.39 4.4 4.24 4.05 0.19 0.15 0.34 0.49 100.36 99.54 0.44 0.43 112 79 695 668 1367 1469 252 149 8.1 4.9 2.6 1.7 17 9.4 23 13 14 10

WA-23 monzqz

ORIGIN OF THE MESOZOIC MAGMATISM 135

3.2 183 14.6 28 127.70 245.00 20.17 79.95 11.62 2.89 8.00 0.90 4.03 0.69 1.80 0.20 1.40 0.21 0.93

Ta V Y Co La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu Eu/Eu*

0.6 157 13.0 17 103.50 197.60 21.66 78.99 10.23 2.64 8.86 1.07 4.52 0.82 2.39 0.30 1.76 0.29 0.86

YM-2 gabbro 0.23 184 13.9 31.4 35.90 74.22 9.11 37.38 6.180 2.040 4.670 0.614 2.928 0.527 1.423 0.170 1.16 0.18 1.18

1.2 132 1 16 57.82 101.80 7.49 27.57 4.06 1.09 3.34 0.53 2.37 0.43 1.14 0.17 1.06 0.17 0.92

WA5 DH-7 monzo dio enclave 0.6 109 12.42 16 40.19 77.49 6.48 29.19 4.79 1.47 4.33 0.6 3.02 0.50 1.41 0.20 1.24 0.20 1.00

YM-5 enclave 0.7 133 12.14 18 48.89 97.18 8.24 35.64 5.86 1.51 4.58 0.63 3.02 0.53 1.44 0.170 1.14 0.180 0.90

ZG-1 monz 1.4 82 14.91 11 56.98 109.7 9.42 38.67 6.51 1.7 5.4 0.72 3.28 0.64 1.73 0.22 1.41 0.23 0.89

DH-33 monz 0.9 64 10 10 39.72 81.60 7.05 30.18 5 1.41 4.13 0.5 2.57 0.43 1.16 0.16 0.94 0.16 0.96

WA-35 monz 1 89 9.51 13 47.08 89.18 7.53 29.8 4.75 1.19 3.56 0.53 2.6 0.43 1.21 0.17 0.91 0.14 0.90

YM-9 monz 1.1 55 8.41 9.1 44.99 79.09 6.72 25.58 3.8 1.16 3.28 0.45 2.15 0.38 0.96 0.12 0.86 0.14 1.02

WA-16 qz. monzqz

LOI Loss on ignition. monz. dio, monzodiorite; qz monz, quartz monzonite; monz, monzonite. Data sources; Chen & Zhai (2003); Chen et al. (2003b).

XG-4 gabbro

Sample no: Rock type:

Table 2. Continued

0.6 37 8.05 4 49.07 89.74 7.55 29.55 4.29 1.11 3.34 0.45 2.08 0.35 0.96 0.12 0.79 0.110 0.91

WA-7 monzqz

1.1 67 12.3 9.6 57.87 107.8 8.98 34.13 5.36 1.34 4.16 0.55 2.96 0.55 1.48 0.2 1.27 0.21 0.88

WA-23 monzqz

0,7 52 8.16 7.2 47.09 86.96 7.03 27.50 4.35 1.08 3.19 0.46 2.02 0.37 0.95 0.120 0.76 0.120 0.90

DH-9 monzqz

1 38 5.73 5.8 35.41 61.47 4.88 18.79 3.74 0.91 2.64 0.35 1.69 0.28 0.67 0.11 0.64 0.11 0.90

DH-22 monz

136 B. CHEN ET AL.

LL-06 dolerite

50.8 0.88 14.86 9.00 0.20 9.73 9.08 2.14 2.78 0.41 0.1 100 0.68 66 1079 1866 167 3.99 1.4 7.5 4.2 10 4.75

Sample no: Rock type:

SiO2 TiO2 Al2O3 FeO MnO MgO CaO Na2O K2O P2O5 LOI Total Mg-no. (ppm) Rb Sr Ba Zr Hf U Th Pb Nb Ta

51.20 0.94 16.00 9.88 0.17 7.64 9.15 2.60 1.59 0.16 0.58 99.893 0.61 56 1023 1094 140 3.14 0.8 3.9 1.8 5.8 3.14

JQ-03 dolerite 53.65 0.97 18.38 7.83 0.16 3.21 6.22 4.8 3.22 0.59 0.8 99.83 0.45 65.2 991 1900 219 5.09 2 4.9 20.2 15.8

R071 enclave 55.69 0.81 18.72 6.44 0.12 2.55 4.74 4.71 4.47 0.54 0.5 99.29 0.44 61.3 953 3285 344 7.22 0.86 4.5 25.1 12.5

R072 enclave 58.64 0.65 12.92 7.21 0.19 8.82 6.10 3.06 2.28 0.20 0.10 100.17 0.71 138 1071 762 212 5.55 3.1 9.1 1.6 16 2.91

XC-01 monz 59.64 0.67 12.97 7.13 0.21 8.44 5.48 2.81 2.47 0.19 0.10 100.12 0.70 165 1044 767 201 5 2.6 9.4 1.9 16 1.25

XC-02 monz

Table 3. Chemical data for the Mesozoic rocks from Jiaodong, North China

61.19 0.65 15.46 5.36 0.10 4.53 5.22 3.86 3.11 0.26 0.10 99.84 0.63 88 1517 4381 301 7.16 3 13.6 23.7 12 1.68

XC-04 monz 58.50 0.67 13.57 6.77 0.15 8.87 5.78 2.86 2.45 0.18 0.10 99.90 0.72 82 1097 1381 200 5.61 3.2 12.5 24.2 17 3.38

XC-09 monz 60.4 0.77 16.22 6.01 0.1 3.06 5 3.98 3.06 0.26 0.5 99.36 0.50 45.7 682 1827 131 3.59 0.37 3.97 10.7 0.66

9.73 0.69

L301 monzqz

67.3 0.46 15.68 3.06 0.05 1.2 2.56 4.54 3.86 0.16 0.21 99.08 0.44 60.3 680 1863 191 5.27 1.13 7.39

L306 qz monzqz

11.7 0.68

57.38 0.86 16.97 7.58 0.12 3.52 6.2 3.87 2.43 0.3 0.34 99.57 0.48 31.9 635 1247 126 3.35 0.34 3

L302 monzqz

10.7 0.55

56.71 0.86 17.01 7.48 0.12 3.57 6.16 3.88 2.16 0.31 1.48 99.74 0.49 33.5 646 1226 132 3.45 0.38 4.33

L63 monzqz

(Continued)

11.4 0.84

65.13 0.53 15.68 3.99 0.06 1.86 3.3 4.31 3.92 0.2 0.8 99.78 0.48 64.8 655 1909 158 4.39 1.31 7.31

L323 monz

ORIGIN OF THE MESOZOIC MAGMATISM 137

20.0 37 48.70 99.8 11.50 45.60 7.84 2.73 7.10 0.83 3.85 0.74 2.05 0.27 1.90 0.27

LL-06 dolerite

20.0 48 35.70 76.0 8.92 35.80 6.56 2.16 6.09 0.80 4.29 0.84 2.30 0.32 1.56 0.27

JQ-03 dolerite 100 21.4 60.00 120.00 13.40 51.00 8.84 2.66 5.87 0.83 4.36 0.86 2.11 0.29 1.84 0.29

65.90 133.00 15.90 61.10 10.40 2.52 7.45 1.02 5.15 1.04 2.62 0.35 2.06 0.33

R072 enclave

119 25.3

R071 enclave 18 41 26.30 50.20 5.87 23.50 5.05 1.49 5.26 0.61 3.31 0.65 1.91 0.240 1.66 0.260

XC-01 monz 17 42 29.9 51.5 5.59 23.9 4.46 1.59 5.02 0.59 3.26 0.64 1.97 0.31 1.75 0.25

XC-02 monz 17 28 74.6 143.00 16.2 58.6 8.74 3.22 7.41 0.78 3.45 0.65 1.82 0.26 1.42 0.21

XC-04 monz

monz monzonite; qz monz, quartz monzonite. Data sources: Mang et al. (2003); Yang et al. (2004); Hu et al. (2005).

V Y Co La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu

Sample no: Rock type:

Table 3. Continued

20 41 37.7 76.2 8.55 31.7 6.01 2.24 6.36 0.76 4.12 0.76 2.14 0.28 2.01 0.27

XC-09 monz 44.2 9.83 6.53 55.5 90.2 9.42 33 4.99 1.42 3.71 0.43 2.26 0.41 1.15 0.15 0.99 0.15

L306 qz monzqz 104 18.3 18.4 42.90 84.50 9.97 39.70 7.06 2.14 6.27 0.79 4.34 0.81 2.21 0.30 1.85 0.270

L301 monzqz

156 25.1 20.5 41.8 85.7 10.9 45.1 8.51 2.48 7.73 1.02 5.75 1.1 3.07 0.42 2.58 0.38

L302 monzqz

141 22.9 19.6 63.20 119.0 13.20 50.0 8.20 2.39 7.28 0.96 5.27 0.98 2.77 0.38 2.29 0.34

L63 monzqz

69.9 13.6 10.5 44.80 77.90 9.59 35.5 5.87 1.59 4.33 0.56 3.04 0.57 1.64 0.22 1.46 0.23

L323 monz

138 B. CHEN ET AL.

98.0

32.0

63,27

25.60

32.90

69.55

51.77 1.35 16.5 8.94 0.2 5.62 7.19 3.16 4.06 0.55 1.1 100.44 0.55 40 1071 1796 299 52.00

49.63 1.22 17.98 10.78 0.15 4.33 7.69 3.28 2.76 0.71 1.57 100.10 0.44 35 1092 1444 261 46.00

SiO2 TiO2 Al2O3 FeO4 MnO MgO CaO Na2O K2O P2O5 LOI Total Mg-no Rb Sr Ba Zr Nb Th Y Pb V Cr Co Ni La

D1131 gabbro

D45 gabbro

Sample no: Rock type:

58.92

91.0

17.40

54.52 1.41 16.61 7.34 0.14 4.62 7.32 3.14 2.92 0.46 1.6 100.08 0.55 50 992 1865 323 47.00

D15 diorite

77.81

69.0

21.40

62.49 0.7 16.41 4.63 0.14 1.96 3.52 3.5 5.11 0.35 0.96 99.77 0.46 121 738 1775 351 29.00

D12 monz

57.25

52.0

17.00

56.87 0.98 19.17 6.37 0.13 2.29 5.32 3.65 4.03 0.46 0.58 99.85 0.42 101 1184 3343 526 40.00

D1360 monz.dio 60 0.84 17.57 5.39 0.07 2.28 4.21 4.25 4.6 0.39 0.53 100.13 0.46 100 945 2151 358 14.10 11.30 18.90 22.10 89.4 26.8 13.2 17.9 67.70

D62 monz

Table 4. Chemical data for the Mesozoic rocks from Dabie, Central China

54.35

44.0

19.30

62.64 0.74 16.91 5.46 0.1 2.05 4.53 3.48 3.11 0.38 0.49 99.89 0.43 100 761 1554 264 29.00

D1317 monz

43.73

76.0

18.70

63.22 0.64 15.65 5.06 0.13 2.25 4.5 3.46 4.16 0.33 0.73 100.13 0.47 100 664 1562 314 26.00

D1464 monz 65.11 0.54 16.94 3.39 0.05 0.92 2.11 4.35 5.75 0.19 0.68 100.03 0.35 130 471 1660 295 16.00 13.20 16.30 22.30 35.2 2.0 3.9 4.1 69.50

D204 monz 66.70 0.39 17.10 2.57 0.03 0.62 1.94 4.25 5.98 0.12 0.35 100.05 0.32 120 418 1597 253 13.40 9.47 16.70 22.90 25.1 5.0 3.6 3.7 48.50

D47 monz

65.59

55.0

15.60

67.26 0.50 15.27 3.95 0.09 1.33 1.90 3.49 4.71 0.29 1.17 99.96 0.40 135 461 1482 238 26.00

90.28

30.0

26.50

71.37 0.37 14.73 2.54 0.07 0.48 1.54 3.16 4.66 0.20 0.60 99.72 0.27 159 230 695 191 17.00

D1362 D1286 qz-monz monz.gr

76.57

137.0 75 23.5

23.90

54.57 1.10 19.30 7.91 0.15 3.87 6.68 3.65 3.21 0.58 0.20 101.22 0.49 65 1200 2604 343 56.00

D50 enclave

(Continued)

72.74 0.23 14.25 1.80 0.020 0.28 1.14 3.52 5.27 0.23 0.06 99.54 0.24 263 239 1036 227 17.10 34.20 16.80 39.70 14.8 2.5 2.1 1.7 84.30

D1172 granite

ORIGIN OF THE MESOZOIC MAGMATISM 139

135.5 18.15 71.85 13.87 2.99 11.68 1.5 7.2 1.49 3.67 0.54 2.85 0.390 0.73

Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu Eu/Eu*

117 15.01 53.84 10.44 2.64 8.52 1.13 5.54 1.14 2.8 0.4 2.1 0.320 0.87

D1131 gabbro 103.3 13.33 44.79 8.53 2.49 7.71 0.91 3.93 0.83 2 0.29 1.62 0.230 0.95

D15 diorite 139.2 18.32 57.93 10.23 2.47 7.18 0.98 4.58 0.97 2.36 0.38 2.19 0.350 0.89

D12 monz 98.77 13.27 45.93 8.28 2.73 6.38 0.82 3.8 0.83 2.19 0.31 1.58 0.240 1.16

D1360 monz.dio 124.00 13.80 53.40 8.16 2.710 5.97 0.757 4.08 0.672 1.97 0.278 1.64 0.267 1.20

D62 monz 91.14 12.13 40.56 7.72 2.000 5.85 0.79 3.92 0.82 2.06 0.31 1.73 0.280 0.92

D1317 monz 75.59 10.06 34.53 6.82 1.670 5.44 0.74 3.72 0.81 2.13 0.3 1.8 0.310 0.85

D1464 monz

Mon monzonite; monz. gr monzogranite; monz. dio, monzodiotrite. Data sources: Jahn et al. (1999); Chen et al. (2002).

D45 gabbro

Sample no: Rock type:

Table 4. Continued

132.00 12.50 40.40 6.04 1.710 4.35 0.61 3.21 0.52 1.41 0.21 1.42 0.191 1.03

D204 monz 93.50 10.00 33.30 5.38 1.490 3.88 0.58 2.98 0.55 1.43 0.227 1.45 0.212 1.01

D47 monz

104.20 152.70 13.50 18.53 40.10 55.08 6.79 9.67 1.680 1.540 4.96 7.30 0.66 1.02 3.14 5.37 0.68 1.18 1.78 3.02 0.280 0.47 1.51 2.65 0.220 0.400 0.90 0.57

D1362 D1286 qz-monz monz.gr

139.20 17.53 60.46 11.27 2.800 8.48 1.11 5.16 1.10 2.75 0.410 2.23 0.340 0.89

D50 enclave

141.00 13.80 43.40 5.92 1.180 4.28 0.599 3.32 0.54 1.48 0.25 1.60 0.254 0.73

D1172 granite

140 B. CHEN ET AL.

ORIGIN OF THE MESOZOIC MAGMATISM

141

Fig. 4. Total alkalis v. silica diagram (Middlemost 1994), showing that the Mesozoic igneous rocks in the NCC are dominated by monzonites and quartz monzonites, with subordinate mafic rocks, syenites and granitoids. S , V, Plutonic and volcanic rocks from Taihang–Yanshan (Li et al. 2001; Liu et al. 2002; Qian et al. 2002; Chen & Zhai 2003; Chen et al. 2003b; Gao et al. 2004), respectively. A, B, Pluntonic and volcanic rocks from Jiaodong (Sun et al. 1996; Qiu et al. 1997; Meng et al. 2003; Yang et al. 2003, 2004; Guo et al. 2004; Hu et al. 2005), respectively. O, Plutonic rocks from Dabie (Ma et al. 1998; Jahn et al. 1999; Wang, et al. 2001; Chen et al. 2002; Pan et al. 2001); no data for volcanic rocks from Dabie are available. s is the Rittmann index defined as (K2O þ Na2O)2/(SiO2 2 43).

in terms of rock types, which include dominant monzonitic to quartz monzonitic rocks (65%) and subordinate mafic rocks (monzogabbroic to monzodioritic; 15%), granitoids (10%) and syenites (10%). These plutonic rocks are accompanied by equally voluminous, contemporaneous volcanic rocks show high-K calc-alkaline to shoshonitic affinity, with 46 –76% SiO2 and 5– 9% total alkalis (Qiu et al. 1997; Li et al. 2001; Liu et al. 2002; Qian et al. 2002; Chen et al. 2003b, 2005). The main constituent minerals are plagioclase, hornblende, pyroxene, quartz, and K-feldspar, with accessory magnetite, titanite, zircon, allanite and apatite. Olivine is occasionally seen in basic members. A systematic decrease in hornblende and pyroxene abundance is coupled with increasing quartz and feldspar from monzogabbro diorites, through monzonites and quartz monzonites, to monzogranites. To understand the petrogenesis of the NCC Mesozoic magmas, we have constructed Harkerstyle diagrams for selected major (Fig. 5) and trace elements (Fig. 6). It is apparent from Figure 5 that CaO, TiO2 and MgO decrease sharply with increasing SiO2, and the data points define linear, continuous trends in these diagrams.

Similarly, we see linear, negatively correlated variation trends between SiO2 and compatible V and Co (Fig. 5). The linear trends are consistent with a magma mixing model for the generation of the Mesozoic igneous rocks. Sr, however, shows scatter of data points in Figure 6, probably as a result of the combined effects of magma mixing and fractionation during magma evolution (to be discussed below), but roughly negative variation trends between SiO2 and Sr can still be seen. It should be noted that most rocks from the NCC have high Sr (500 –2000 ppm) and Ba (700 – 3500 ppm) abundances (Tables 2–4). Mafic enclaves show chemical compositions intermediate between the mafic and felsic rocks (Tables 2–4; Figs 5 and 6), suggesting that they are also mixtures between mafic and felsic magmas. In the chondrite-normalized rare earth element (REE) patterns (Fig. 7), the Mesozoic rocks from the NCC are all characterized by similarly highly fractionated REE patterns with significant depletion of heavy REE (HREE) and enrichment of light REE (LREE), and particularly with minor anomalies of Eu. Moreover, the intermediate to felsic rocks show REE patterns similar to those of the coeval

142

B. CHEN ET AL.

Fig. 5. Plots of SiO2 v. CaO, TiO2 and MgO for igneous rocks from Taihang–Yanshan, Jiaodong and Dabie. ( S , A, W, Mafic enclaves from Taihang– Yanshan, Jiaodong and Dabie, respectively. Data sources are the same as in Figure 4.

mafic rocks. This again suggests that the intermediate to felsic rocks could be genetically linked in some way to the mantle-derived mafic rocks. Mafic enclaves have REE patterns similar to those of both their host rocks and mafic plutons (shown as coarse grey lines in Fig. 7). The features of low HREE and high Sr are further illustrated in Figure 8, in which pure crustal melts (granitic in composition; Jung et al. 2003) are also shown for comparison. It can be seen from Figure 8 that the NCC Mesozoic magmas are characterized by high Sr/Y and low Y (HREE), in contrast to the low Sr/Y and high Y signatures of pure crustal melts.

Nd – Sr isotopic data Figure 9 shows a compilation of published Nd –Sr isotopic data (recalculated at t ¼ 130 Ma) for the Mesozoic rocks from the NCC. Initial Nd and Sr

isotopic compositions for the mafic rocks from the Taihang–Yanshan orogen are 1Nd(t) ¼ 28 to 214 and ISr ¼ 0.7054–0.706, and for most coeval intermediate to felsic rocks 1Nd(t) ¼ 212 to 218 and ISr ¼ 0.7055– 0.707 (Chen & Zhai 2003; Chen et al. 2003a, 2004). As seen in Figure 9, data points for the intermediate to felsic rocks from Taihang– Yanshan plot between the mafic rocks and lower continental crust (LCC), which is compatible with the model for the genesis of the intermediate to felsic rocks by mixing processes between mantle-derived mafic and LCC-derived granitic magmas. Compared with the mafic rocks from Taihang–Yanshan, the mafic rocks from Dabie and Jiaodong show significantly lower 1Nd(t) values and higher ISr ¼ ratios, with 1Nd(t) ¼ 29 to –16, ISr ¼ 0.7065–0.708 and 1Nd(t) ¼ 29 to –16, ISr ¼ 0.7075–0.710, respectively. Apparently, the mantle source for the mafic rocks from Taihang–Yanshan differs from that for mafic rocks from Dabie and Jiaodong in terms of

ORIGIN OF THE MESOZOIC MAGMATISM

143

Fig. 6. Plots of SiO2 v. Co, V and Sr for igneous rocks from Taihang– Yanshan, Jiaodong and Dabie. ( S , A, W , Mafic enclaves from Taihang– Yanshan, Jiaodong and Dabie, respectively. Data sources are the same as in Figure 4.

isotopic compositions; the reason will be discussed below. Similarly, intermediate to felsic rocks from Dabie and Jiaodong also plot between related mafic rocks and LCC, which is in agreement with the magma mixing model suggested for their genesis. Figure 10 is constructed to delve more deeply into the source characteristics of the Mesozoic magmas from the NCC. Also shown for comparison are the fields of LCC (Jahn et al. 1987), subcontinental lithospheric mantle (SCLM; Chen & Zhai 2003; Chen et al. 2004) and depleted mantle. It is apparent that data points for the Mesozoic magmas from all three regions (Taihang– Yanshan, Jiaodong and Dabie) of the NCC are confined between the SCLM and LCC. Mafic rocks show higher 1Nd(t) values than related intermediate to felsic rocks, although overlaps of data points exist between the two. It appears that material contribution from depleted mantle is negligible. This also suggests that the Mesozoic magmas of the

NCC are mixtures between SCLM-derived mafic magmas and LCC-derived granitic melts.

Discussion and conclusions Origin of the Mesozoic rocks of the NCC As shown above, the Mesozoic rocks of the eastern NCC range in composition from gabbroic to monzogranitic, and show high-K calc-alkaline to shoshonitic affinity, high Sr–Ba abundance and Sr/Y ratios, highly fractionated REE patterns with minor Eu anomalies, and highly enriched Sr –Nd isotopic compositions. Many papers have been published (e.g. Zhang et al. 2001; Chen et al. 2002, 2003a, b, 2004; Liu et al. 2002; Qian et al. 2002; Xu et al. 2002) in the past decade to interpret these grochemical characteristics and origin of the Mesozoic rocks, but it remains an issue of controversy. Most workers advocate that the Mesozoic

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Fig. 8. Plot of Sr/Y v. Y for the Mesozoic rocks from Taihang– Yanshan, Dabie and Jiaodong. Data points for the mafic rocks (SiO2 ,52%) from the NCC are not shown, but their field is indicated by a dashed line. Also shown for comparison is the field for pure crustal melts (granitic) from partial melting of TTG-like gneisses (Jung et al. 2003). Modelling results show how much fractionation (10% graduation) is needed to achieve the spread of data points (see text for further explanation), based on Rayleigh’s law. The parent magma is represented by the open star in the figure, and is assumed to have 13 ppm Y, 900 ppm Sr and Sr/Y ¼ 69. Trend 1 and trend 2 represent residual melts after fractionation of dominantly ferromagnesian phases (hornblende 68% þ Cpx 26% þ plagioclase 5% þ apatite 1%) and fractionation of combined ferromagnesian phases and plagioclase (hornblende 26% þ Cpx 12% þ plagioclase 58% þ apatite 4%), respectively. Distribution coefficients of Y and Sr used in the modelling are from Rollinson (1993). Data sources are the same as in Figure 4.

Fig. 7. Chondrite-normalized REE patterns for the Mesozoic rocks from Taihang–Yanshan (data from Table 2), Jiaodong (data from Table 3) and Dabie (data from Table 4). It should be noted that intermediate to felsic rocks show REE patterns similar to those of the related mafic rocks from Taihang–Yanshan (samples XG-4, YM-2 and WA-5), Jiaodong (samples LL06 and JQ03) and Dabie (samples D45 and D1131), and also similar to those of mafic enclaves (shown as wide grey lines). They all are characterized by highly enriched LREE and depleted HREE and minor Eu anomalies. Chondrite values used in normalization are from Masuda et al. (1973).

intermediate to felsic rocks originated from melting of mafic lower crust that was either (1) beneath a thickened crust (e.g. Zhang et al. 2001; Liu et al. 2002), or (2) delaminated to mantle depths after being transformed into eclogites (e.g. Xu et al. 2002; Gao et al. 2004). Both models emphasize that partial melting of mafic lower crust occurred at high pressures leaving behind eclogitic residues in the source. As a consequence, partial melts in equilibrium with eclogitic residues could have high Sr/Y, low Y and highly differentiated REE patterns (Zhang et al. 2001; Xu et al. 2002). However, the two models have difficulties explaining some other petrological and geochemical signatures. It is apparent from Figures 4 and 5 that the Mesozoic magmas show a wide range of chemical composition ranging from mafic to granitic, and dominantly are monzonitic with relatively low silica contents (55–64%). However, partial melting of basaltic LCC generally produces melts of tonalitic, granodioritic and granitic composition

ORIGIN OF THE MESOZOIC MAGMATISM

Fig. 9. Plot of 1Nd(t) v. initial 87Sr/86Sr for rocks from Taihang–Yanshan, Dabie and Jiaodong (data sources as in Fig. 4). LCC, lower continental crust; UCC, upper continental crust (Jahn et al. 1987). SCLM, subcontinental lithospheric mantle of the NCC (Chen et al. 2003a; Chen & Zhai 2003). Also shown are isotopic modeling results (10% graduation) using a simple mixing model (Langmuir et al. 1978) between SCLM (ISr ¼ 0.7055, 1Nd(t) 28.5, Sr 100 ppm, Nd 8.3 ppm; Chen et al. 2003a; Chen & Zhai 2003) and LCC (ISr ¼ 0.709, 1Nd(t) ¼ 230, Sr 300 ppm, Nd 24 ppm; Chen & Jahn 1998) and UCC (ISr ¼ 0.719, 1Nd(t) 212, Sr 350 ppm, Nd 25 ppm; Chen & Jahn 1998) of the Yangtze Block, respectively.

at reasonable degrees of partial melting (Rushmer 1991), with a relatively high, narrow range of silica contents. This is not the case for the Mesozoic magmas from the NCC. Moreover, regardless of the degree of partial melting, such melts tend to have low Mg-number (¼molar Mg/(Mg þ 0.9  Fe)) according to experimental data (Wolf & Wyllie 1994; Rapp & Watson 1995), which does not agree with the typically high Mg-number of the Mesozoic magmas in the NCC (Chen et al. 2002, 2003a, b, 2004; Liu et al. 2002; Qian et al. 2002;

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Xu et al. 2002; Yang et al. 2004). Figure 11 shows the high Mg-number characteristics of the Mesozoic magmas. Also shown for comparison are experimental data from partial melting of basalts (Sen & Dunn 1994; Rapp & Watson 1995) and pure crustal melts (granitic) from reworking of basement gneisses Jung et al. 2003). Bulk composition of the mafic starting material used in the experiments is alkali-rich basalt, high-alumina basalt and low-potassium tholeiite (Sen & Dunn 1994; Rapp Watson 1995). The partial melting experiments were mostly conducted at pressures of 8–30 kbar and temperatures of 900–1100 8C, with water provided solely via amphibole breakdown. These conditions could be achieved by significant undeplating of basalts at the base of continental crust, where crust thickening is a consequence of underplating. We note from Figure 11 that the Mesozoic magmas of Taihang–Yanshan, Jiaodong and Dabie are basically inseparable; they show significant variation in chemical composition ranging continuously from gabbroic to granitic. Apparently, they have significantly higher Mg-number (mostly .0.4) than the experimental melts from partial melting of basalts, suggesting that the NCC Mesozoic magmas could not be derived from melting of mafic lower crust alone, and mantle input is thus indicated. This is compatible with the model of magma mixing. Model (1) is thus rejected. Model (2) is an improved version of model (1), which was proposed by Xu et al. (2002) and Gao et al. (2004) to explain the high Mg-number of the Mesozoic rocks of the NCC by suggesting that mafic lower continental crust was first transformed into eclogites under overthickened crust and subsequently delaminated to mantle depths to undergo partial melting, and that the Mg-number of the resultant partial melts consequently was raised via interaction with mantle rocks (Rapp et al. 1999) en route to crustal levels. Although

Fig. 10. Plots of 1Nd(t) v. intrusive age for the Mesozoic rocks from Taihang– Yanshan, Jiaodong and Dabie. Also shown for comparison are the fields for the LCC (Jahn et al. 1987), SCLM (Chen et al. 2003a; Chen & Zhai 2003) and depleted mantle.

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Fig. 11. Plot of Mg-number v. SiO2 for the Mesozoic rocks from the NCC. Data sources are the same as in Figure 4. Also shown for comparison are experimental data from melting of basaltic rocks (grey diamond Sen & Dunn 1994; Rapp & Watson 1995) and pure crustal melts (granitic) from melting of basement gneisses (Jung et al. 2003). S, Taihang– Yanshan; A, Jiaodong; W, Dabie.

this model can reasonably explain the elevated Mg-number, it is inconsistent with some other observations as described below. We see no signs of a crustal-thickening mechanism in the eastern NCC in Mesozoic times, which would be something like that (related to subduction of the Indian plate beneath the Euro-Asian continent) in the Tibetan plateau (e.g. Lombardo & Rolfo 2000). On the contrary, structural and basin analysis (e.g. Zhao et al. 2004) all indicate a phase of intense extension starting from 160 Ma to 110 Ma in the eastern NCC, during which the majority of the Mesozoic magmas in the NCC were generated. That is, the Mesozoic rocks were emplaced in an extensional regime, and no overthickened crust could be expected during that time. Therefore, mafic lower crust is less likely to be eclogitized at that time, and thus large-scale delamination of mafic lower crust into mantle depths is less likely to happen without formation of voluminous eclogites in the lower crust. This is supported by the fact that few eclogitic enclaves were found in the Mesozoic plutons. Therefore, model (2) is not favoured, though it cannot be completely ruled out. We favour a mixing model between mantle-derived mafic magmas and granitic melts of crustal origin in an extensional regime for the genesis of the NCC Mesozoic rocks, as discussed below. Mafic magmas from the NCC are characterized by highly fractionated REE patterns with minor Eu anomalies, depleted HREE (and Y) and much enriched LREE (Fig. 7). They have high large ion lithophile element (LILE) abundances, such as Sr up to 2000 ppm (Fig. 6), and consequently have

high Sr/Y ratios (Fig. 8). Moreover, these mafic rocks have higher MgO (Fig. 5) and Mg-number (Fig. 11) than the intermediate to felsic rocks. Pure crustal melts (granitic in composition) derived from partial melting of TTG-like gneisses normally have higher Y (and HREE), lower Sr and thus lower Sr/Y ratios than the Mesozoic rocks from the NCC (Fig. 8), and have very depleted Mg-numbers (Fig. 11). This leads us to suggest that the Mesozoic intermediate to felsic rocks in the NCC could form by mixing between the coeval mafic magma and granitic melts derived from melting of old lower crustal rocks (dominantly Archaean TTG gneisses; Jahn et al. 1987). That is, the overall geochemical signatures such as high Sr–Ba, highly depleted HREE and minor Eu anomalies, and, particularly, high Mg-numbers of the Mesozoic intermediate to felsic rocks are ultimately traceable to the mantlederived mafic magmas, as suggested by Fowler & Henney (1996) for many other high Sr –Ba granitoids. As seen in te chondrite-normalized REE patterns (Fig. 7), the intermediate to felsic rocks basically show REE patterns similar to those of the coeval mafic rocks and microgranular mafic enclaves, which we interpret as inherited from mixing of the mantle-derived mafic magmas. The magma mixing model is strongly supported by our petrological studies. As stated above, microgranular mafic enclaves can frequently be seen within the intermediate to felsic plutons of the NCC (Fig. 2a–c), which are coeval with their host rocks. Moreover, compositional and textural disequilibrium (resorbed An-rich cores mantled by abrupt Ab-rich rims; Fig. 3a and b; Table 1) can easily be found in plagioclase grains from both mafic enclaves and host rocks. These observations are typically viewed as a strong indication of magma mixing between mafic and granitic magmas (e.g. Benito et al. 1999; Janousek et al. 2004; Kemp 2004). This is further supported by the linear and continuous variation trends of SiO2 v. CaO, TiO2 and MgO (Fig. 5), and SiO2 v. Co, V and, to a less extent, Sr (Fig. 6). In addition, the magma mixing model is also supported by Nd–Sr isotopic data. As seen in Figures 9 and 10, mafic rocks plot between the field of SCLM and LCC, and have 1ND values close to, but variably lower than that of SCLM. This probably reflects incorporation of variable proportions of LCC during magma ascent to crustal levels (Chen & Zhai 2003; Chen et al. 2003a, b). Most intermediate to felsic rocks (dominant rock types in the NCC) have Nd (and Sr) isotopic compositions intermediate between mafic rocks and the Precambrian LCC, again implying that they could form by mixing of mantle-derived mafic magmas with LCC-derived granitic melts. Modelling studies based on a simple mixing model

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(Langmuir et al. 1978) indicate that less than 40% of the LCC-derived granitic melts were variably involved in generating the intermediate to felsic magmas through mixing with the mafic magma derived from partial melting of enriched SCLM (Fig. 9). Although magma mixing has played an important role in generating the wide spectrun of rock types of the Mesozoic igneous rocks, processes of crystal fractionation also contributed to modification of the chemistry of the Mesozoic rocks. As shown in Figure 8, mafic rocks have higher Sr/Y ratios and lower Y than pure crustal melts (granitic in composition); mixing between granitic and mafic magmas can produce the Mesozoic rocks whose data points plot continuously between the two. However, many data points do not plot between the fields for mafic rocks and pure crustal melts in Figure 8. We attribute this to fractionation of the hybrid magma. Chen et al. (2002, 2003a, b) suggested that the Mesozoic magmas experienced significant fractionation of dominantly ferromagnesian phases (Cpx and hornblende) and subordinate plagioclase during magma evolution, based on the minor Eu anomalies of the Mesozoic rocks in the REE patterns. This can significantly enhance the Sr/Y, and decrease the abundance of Y (and Yb) in residual melts, particularly when plagioclase removal is minor (see fractionation trend (1) in Fig. 8), because Cpx and hornblende have low distribution coefficients of Sr (DSr , 1.0) and high DY and DYb (.1.0). Fractionation trend (2) in Figure 8 represents a process in which plagioclase removal is equally important, because plagioclase fractionation inevitable decreased Sr abundance and thus raised Sr/Y ratios of the residual melts. To better understand the effects of crystal fractionation on chemistry of the Mesozoic magmas, we have carried out geochemical modeling based on Rayleigh’s law, assuming that the parental hybrid magma (shown as a star in Fig. 8) has 13 ppm Y and Sr/Y ¼ 69. Modelling results suggest that fractionation trends (1) and (2) represent residual melts after variable proportions (10% graduation) of fractionation of hornblende (68%) þ Cpx (26%) þ plagioclase (5%) þ apatite (1%) and of hornblende (26%) þ Cpx (12%) þ plagioclase (58%) þ apatite (4%), respectively. The spread of data points in Figure 8 can reasonably be accounted for by fractionation of dominantly ferromagnesian phases (hornblende þ Cpx), subordinate plagioclase and accessory apatite.

Characteristics of mantle source regions for mafic rocks The uniformly low 1Nd values (28 to –15; Fig. 9), together with the highly enriched LILE and LREE

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abundances (Tables 2– 4; Figs 6 and 7) of the mafic rocks, suggest their derivation from melting of enriched SCLM sources (Chen et al. 2002, 2003a, 2004; Chen & Zhai 2003; Yang et al. 2004). The enriched SCLM sources could be created by interaction of mantle rocks with slabderived fluids during an earlier phase (in late Archaean to early Proterozoic times) of subduction (McCulloch & Gamble 1991; Zhao et al. 2002) and/or with volatile-rich, low-density melts released from the asthenosphere and leaked to the SCLM above (McKenzie 1989). Partial melting of the enriched mantle source can produce melts with high Sr–Ba and LREE, and more so if the degree of partial melting is low (e.g. Rogers et al. 1998). Later stage of fractionation of ferromagnesian phases also contributed to the high-Sr signature as a result of their low distribution coefficients of Sr (DSr). The low Y (and HREE) of the mafic magmas is mainly related to the presence of residual garnet in mantle sources, which is further enhanced by removal of ferromagnesian minerals in laterstage magma evolution. The Taihang– Yanshan mafic rocks show typical EM 1-type enriched mantle signatures with ISr ¼ 0.7053–0.7059 and 1Nd(t) ¼ 28 to 214 (Fig. 9). Mafic rocks from Dabie and Jiaodong, however, show higher ISr ratios and lower 1Nd(t) values than those from the Taihang–Yanshan orogen, data points of the former thus define a rightshift trend in Figure 9. The right-shift trend reflects an isotopic heterogeneity of mantle sources beneath the NCC, and the reason for this remains unclear. Many recent studies (Li et al. 1993; Rowley et al. 1997; Jahn et al. 1999) have indicated that the Yangtze continent was subducted to mantle depths during Triassic collision between the Yangtze Block and NCC, and experienced UHP metamorphism at about 220 Ma. Subsequent exhumation of the subducted continent formed the prominent Dabie –Sulu UHP belt (Fig. 1). Taking into account the fact that most of the Dabie and Jiaodong Mesozoic rocks were emplaced in or close to the UHP belt at around 130 Ma, we suggest that the already enriched SCLM sources beneath the UHP belt (southern part of the NCC) might have been isotopically modified, prior to the Mesozoic magmatism, through interaction with subducted Yangtze continental materials that have significantly higher Sr isotopic ratios (0.718–0.725 for the upper continental crust (UCC) and 0.706–0.712 for the LCC; Chen & Jahn 1998) and lower 1Nd(t) values (210 to 213 for the UCC and around 230 for the LCC; Chen & Jahn 1998) than the EM1-type enriched mantle source beneath the NCC (typically with 87 Sr/86Sr ¼ 0.7053–0.7059 and 1Nd(t) ¼ 28 to 212; Chen & Zhai 2003; Chen et al. 2003a). As a result, the mafic rocks from Dabie and Jiaodong

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have acquired higher Sr isotopic ratios and lower 1Nd(t) values than those from the Taihang – Yanshan orogen, which is far from the UHP belt, and the mantle source beneath which thus escaped being influenced by the subducted Yangtze continental materials. Isotopic modelling using a simple mixing model (Langmuir et al. 1978) suggests that both the UCC and LCC components of the Yangtze continent were variably involved in the processes of interaction with EM1-type mantle material beneath the NCC (Fig. 9). The proportions of UCC components involved in the mantle source are estimated at 4–12%, but those of LCC components involved are hard to estimate because of possible contamination of lower crustal components of the NCC (having 1Nd(t) values as low as to 235; Jahn et al. 1987) during magma ascent. As seen in Figure 9, the isotopic compositions of the mafic rocks from Dabie and Jiaodong show significant variations; their 1Nd(t) and ISr are roughly inversely correlated, and extend towards the LCC field. This suggests that the mantle-derived mafic magmas could have been contaminated by the lower crustal material of the NCC en route to crustal levels. Therefore, the right-shift trend and extension towards the LCC of data points for the Dabie and Jiaodong rocks in a plot of 1Nd(t) v. ISr (Fig. 9) are probably accounted for first by interaction of the EM1-type mantle source with subducted Yangtze continental components in late Triassic time, then by contamination by lower continental crustal components of the NCC during magma evolution in the Mesozoic.

Geodynamic setting and petrogenetic model: a summary What is the cause of the sudden surge of Mesozoic magmatism in the NCC? A popular model holds that the lithospheric destruction was related to the loss of physical integrity of the craton, caused by the Triassic collision between the NCC and Yangtze Block (Menzies et al. 1993; Xu 2001; Gao et al. 2002). The Mesozoic magmatism was thus considered by many (e.g. Mao et al. 2003) as post-collisional magmatism developed in an intracontinental extensional regime. We alternatively suggest that the Mesozoic magmatism was probably related to subduction of the palaeo-Pacific slab beneath East China. The Mesozoic igneous rocks and contemporaneous sedimentary basins in the NCC were dominantly distributed northeasterly, roughly parallel to the subduction zone lying outside the Japan island chain (Maruyama 1997). This model is further supported by the extensive accretion of arc complexes in the eastern margin (Japan island train) of the East Asian continent

during Jurassic times (Maruyama 1997). The voluminous Mesozoic igneous rocks in NE China (Wu et al. 2003) and SE China (Zhou & Li 2000; Fig. 1) were also interpreted to have formed as a result of subduction of the palaeo-Pacific slab, based mainly on the younging trend of magmatism from coastal to inland regions. We thus suggest that the entire Mesozoic magmatism in East China is part of the East Asian continental arc, a model also advocated by Sengor & Natal’in (1996). The above petrological and geochemical data collectively suggest that both the enriched SCLM and Precambrian LCC contributed to the sources of the Mesozoic intermediate to felsic rocks. The genesis of the intermediate to felsic magmas can be described briefly as follows. Subduction of the palaeo-Pacific slab beneath the East Asian continent transformed the eastern part of the NCC into an active continental margin and, consequently, a back-arc extensional setting was developed (Fig. 1). This might have induced the upwelling of asthenosphere, which, in turn, triggered extensive partial melting of the enriched SCLM beneath the NCC, producing voluminous mafic magmas (with high Sr, Ba, LREE and K, and low Y and Yb). Underplating of the mafic magmas in the LCC provided a heat source for partial melting of TTG-dominated LCC basement rocks (John et al. 1987), producing granitic melts, which subsequently mixed with the underplating mafic magmas to form hybrid magmas. The hybrid magmas underwent fractionation of dominantly hornblende, pyroxene and plagioclase during magma evolution (Fig. 8), producing the wide spectrum of intermediate to felsic rocks. Therefore, most mafic magmas were consumed through mixing in the LCC to produce hybrid magmas that were parental to the Mesozoic intermediate to felsic rocks. That is why the volume of mafic rocks is subordinate to that of intermediate to felsic rocks in the field. Minor magmatism is seen to the west of the Taihang– Yanshan orogen (Fig. 1), probably as a result of the cessation of subduction or oceanward back-stepping of the subduction zone. We are grateful to Z. Y. Chu and R. H. Zhang for their assistance in isotopic analyses, and to H. Rollinson and M. Fowler for their constructive comments, which led to substantial improvement of this paper. This study was supported by three NSFC grants (No. 40502009 to W. T. and Nos. 40625005 & 40372033 to B.C.

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source region. Acta Petrologica Sinica, 18, 257–274 [in Chinese with English abstract]. L OMBARDO , B. & R OLFO , F. 2000. Two contrasting eclogite types in Himalayas: implications for the Himalayan orogeny. Journal of Geodynamics, 30, 37–60. M A , C. Q., L I , Z., E HLERS , C., Y ANG , K. & W ANG , R. 1998. A post-collisional magmatic plumbing system: Mesozoic granitoid plutons from the Dabieshan highpressure and ultrahigh-pressure metamorphic zone, east– central China. Lithos, 45, 431– 456. M A , X. 1987. Lithospheric Dynamics Map of China and Adjacent Seas (1:4 000 00), and Explanatory Notes. Geological Publishing House, Beijing. M AO , J. W., W ANG , Y. T., Z HANG , Z. H., Y U , J. J. & N IU , B. G. 2003. Geodynamic settings of Mesozoic large-scale mineralization in North China and adjacent areas. Science in China (D), 46, 838–851. M ARUYAMA , S. 1997. Pacific-type orogeny revised: Miyashiro-type orogeny proposed. Island Arc, 6, 91–120. M ASUDA , A., N AKAMURA , N. & T ANAKA , T. 1973, Fine structures of mutually normalized rare-earth patterns of chondrites. Geochimica et Cosmochimica Acta, 37, 239 –248. M C C ULLOCH , M. T. & G AMBLE , J. A. 1991, Geochemical and geodynamical constraints on subduction zone magmatism. Earth and Planetary Science Letters, 102, 358 –374. M C K ENZIE , D. P. 1989. Some remarks on the movement of small melt fractions in the mantle. Earth and Planetary Science Letters, 95, 53–72. M ENG , F. C., X U , Z. Q., Z HANG , Z. M. & L IU , F. L. 2003. Geochemical characteristics of the Mesozoic postcollisional granites in northern Jiangsu, China and their geological implications. Acta Geologica Sinica, 77, 566 –576. M ENZIES , M. A. & X U , Y. G. 1998. Geodynamics of the North China Craton. In: F LOWER , M. F. J., C HUNG , S. L., L O , C. H. & L EE , T. Y. (eds) Mantle Dynamics and Plate Interactions in East Asia. Geophysical Monograph, American Geophysical Union, 27, 155– 165. M ENZIES , M. A., F AN , W. M. & Z HANG , M. 1993. Palaeozoic and Cenozoic lithoprobes and loss of .120 km of Archean lithosphere, Sino-Korean craton, China. In: P RICHARD , H. M., A LABASTER , T., H ARRIS , N. B. W. & N EARY , C. R. Magmatic Processes Plate Tectonics Geological Society, London, Special Publications, 6, 71– 81. M IDDLEMOST , E. A. K. 1994. Naming material in the magma/igneous rock system. Earth Science Reviews, 37, 215 –224. P AN , G. Q., L U , X. C. & Y U , H. B. 2001. Petrological and geochemical characteristics of Mesozoic adakite from northern Huaiyang and discussion on its genesis. Acta Petrologica Sinica, 17, 541– 500 [in Chinese with English abstract]. Q IAN , Q., C HUNG , S. L., L EE , T. Y. & W EN , D. J. 2002. Geochemical characteristics and petrogenesis of the Badaling high Sr –Ba granitoids: a comparison of igneous rocks from North China and the Dabie–Sulu orogen. Acta Petrologica Sinica, 18, 275 –292.

Q IU , J. S., W ANG , D. Z., Z ENG , J. H. & M CE I NNES , B. I. A. 1997. Study on trace element and Nd–Sr isotopic geochemistry of the Mesozoic K-rich volcanic rocks and lamprophyres in western Shandong. Geological Journal of China Universities, 3, 384– 395 [in Chinese]. R APP , R. P. & W ATSON , E. B. 1995. Dehydration melting of metabasalt at 8 –32 kbar: implications for continental growth and crust–mantle recycling. Journal of Petrology, 36, 891– 931. R APP , R. P., S HIMIZU , N., N ORMAN , M. D. & A PPLEGATE , G. S. 1999. Reaction between slab-derived melts and peridotite in the mantle wedge: experimental constraints at 3.8 GPa. Chemical Geology, 160, 335–356. R OGERS , N. W., J AMES , D., K ELLEY , S. P. & DE M ULDER , M. 1998. The generation of potassic lavas from the eastern Virunga province, Rwanda. Journal of Petrology, 39, 1223–1247. R OLLINSON , H. 1993. Using Geochemical Data: Evaluation, Presentation, Interpretation. Longman, Harlow, 108– 111. R OWLEY , D. B., X UE , R., T UCKER , R. D., P ENG , Z. X., B AKER , J. & D AVIS , A. 1997. Ages of ultrahighpressure metamorphism and protolith orthogneisses from the eastern Dabieshan: U–Pb zircon geochronology. Earth and Planetary Science Letters, 151, 191–203. R USHMER , T. 1991. Partial melting of two amphibolites: contrasting experimental results under fluid-absent conditions. Contributions to Mineralogy and Petrology, 107, 41– 59. S EN , C. & D UNN , T. 1994. Dehydration melting of a basaltic composition amphibolite at 1.5 and 2.0 GPa: implications for the origin of adakites. Contributions to Mineralogy Petrology, 117, 394– 409. S ENGO¨ R , A. M. C. & N ATAL ’ IN , B. A. 1996. Paleotectonics of Asia: fragments of a synthesis, In: Y IN , A. & H ARRISON , T. M. (eds) The Tectonic Evolution of Asia. Cambridge University Press, Cambridge, 486–641. S UN , J. G., L IU , C. H. & Z HENG , C. Q. 1996. Genetic mechanism of the Mesozoic intermediate– acid granitic complex in eastern Jiaodong area. Geological Journal of China Universities, 2, 207–217. W ANG , L. G., Q IU , Y. M., M C N AUGHTON , N. J., G ROVES , D. I., L UO , Z. K., M IAO , L. & L IU , Y. K. 1998. Constraints on crustal evolution and gold metallogeny in the northwestern Jiaodong Peninsula, China, from SHRIMP U– Pb zircon studies of granitoids. Ore Geology Review, 13, 243– 258. W ANG , Q., X U , J. F. & Z HAO , Z. H. 2001. The petrogenesis and geodynamic significance of HREE depleted granitoids during Yanshan period in the Dabie Mountains. Acta Petrologica Sinica, 17, 551–564 [in Chinese with English abstract]. W OLF , M. B. & W YLLIE , J. P. 1994. Dehydration-melting of amphibolite at 10 kbar: the effects of temperature and time. Contributions to Mineralogy and Petrology, 115, 369– 383. X U , J. F., S HINJO , R., D EFANT , M. J., W ANG , Q. & R APP , R. P. 2002. Origin of Mesozoic adakitic intrusive rocks in the Ningzhen area of east China: partial

ORIGIN OF THE MESOZOIC MAGMATISM melting of delaminated lower continental crust. Geology, 30, 1111– 1114. X U , Y. G. 2001. Thermo-tectonic destruction of the Archean lithospheric keel beneath the Sino-Korean craton in China: evidence, timing and mechanism. Physics and Chemistry of the Earth (A), 26, 747– 757. Y ANG , J. H., C HU , M. F., L IU , W. & Z HAI , M. G. 2003. Geochemistry and petrogenesis of Guojialing granodiorites from the northwestern Jiaodong peninsula, eastern China. Acta Petrologica Sinica, 19, 692– 700 [in Chinese with English abstract]. Y ANG , J. H., C HUNG , S. L., Z HAI , M. G. & Z HOU , X. H. 2004. Geochemical and Sr –Nd –Pb isotopic compositions of mafic dikes from the Jiaodong Peninsula, China: evidence for vein-plus-peridotite melting in the lithospheric mantle. Lithos, 73, 145– 160. Z HANG , Q., Q IN , K., W ANG , Y., Z HANG , F. Q. & L IU , H. T. 2001. Preliminary study on the components of the lower crust in east China plateau during Yanshanian period: constraints from Sr and Nd isotopic compositions of adakite-like rocks. Acta Petrologica Sinica, 17, 505– 513 [in Chinese with English abstract].

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Z HAO , G. C., W ILDE , S. A. & C AWOOD , P. A. 2002. SHRIMP U– Pb zircon ages of the Fuping complex: implications for Late Archean to Paleoproterozoic accretion and assembly of the North China Craton. American Journal of Science, 302, 191– 226. Z HAO , Y., Z HANG , S. H., X U , G., Y ANG , Z. Y. & H U , J. M. 2004. The Jurassic major tectonic events of the Yanshanian intraplate deformation belt. Geological Bulletin of China, 23, 854– 863. Z HENG , J. P., O’R EILLY , S. Y., G RIFFIN , W. L., L U , F., Z HANG , M. & P EARSON , N. J. 2001. Relict refractory mantle beneath the eastern North China block: significance for lithosphere evolution. Lithos, 57, 43–66. Z HOU , X. H., S UN , M., Z HANG , G. & C HEN , S. H. 2002. Continental crust and lithospheric mantle beneath North China: isotopic evidence from granulite xenoliths in Hannuoba, Sino-Korean craton. Lithos, 62, 111– 124. Z HOU , X. M. & L I , W. X. 2000. Origin of late Mesozoic igneous rocks in southeastern China: implications for lithosphere subduction and underplating of mafic magmas. Tectonophysics, 326, 269–287.

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source region. Acta Petrologica Sinica, 18, 257–274 [in Chinese with English abstract]. L OMBARDO , B. & R OLFO , F. 2000. Two contrasting eclogite types in Himalayas: implications for the Himalayan orogeny. Journal of Geodynamics, 30, 37–60. M A , C. Q., L I , Z., E HLERS , C., Y ANG , K. & W ANG , R. 1998. A post-collisional magmatic plumbing system: Mesozoic granitoid plutons from the Dabieshan highpressure and ultrahigh-pressure metamorphic zone, east– central China. Lithos, 45, 431– 456. M A , X. 1987. Lithospheric Dynamics Map of China and Adjacent Seas (1:4 000 00), and Explanatory Notes. Geological Publishing House, Beijing. M AO , J. W., W ANG , Y. T., Z HANG , Z. H., Y U , J. J. & N IU , B. G. 2003. Geodynamic settings of Mesozoic large-scale mineralization in North China and adjacent areas. Science in China (D), 46, 838–851. M ARUYAMA , S. 1997. Pacific-type orogeny revised: Miyashiro-type orogeny proposed. Island Arc, 6, 91–120. M ASUDA , A., N AKAMURA , N. & T ANAKA , T. 1973, Fine structures of mutually normalized rare-earth patterns of chondrites. Geochimica et Cosmochimica Acta, 37, 239 –248. M C C ULLOCH , M. T. & G AMBLE , J. A. 1991, Geochemical and geodynamical constraints on subduction zone magmatism. Earth and Planetary Science Letters, 102, 358 –374. M C K ENZIE , D. P. 1989. Some remarks on the movement of small melt fractions in the mantle. Earth and Planetary Science Letters, 95, 53–72. M ENG , F. C., X U , Z. Q., Z HANG , Z. M. & L IU , F. L. 2003. Geochemical characteristics of the Mesozoic postcollisional granites in northern Jiangsu, China and their geological implications. Acta Geologica Sinica, 77, 566 –576. M ENZIES , M. A. & X U , Y. G. 1998. Geodynamics of the North China Craton. In: F LOWER , M. F. J., C HUNG , S. L., L O , C. H. & L EE , T. Y. (eds) Mantle Dynamics and Plate Interactions in East Asia. Geophysical Monograph, American Geophysical Union, 27, 155– 165. M ENZIES , M. A., F AN , W. M. & Z HANG , M. 1993. Palaeozoic and Cenozoic lithoprobes and loss of .120 km of Archean lithosphere, Sino-Korean craton, China. In: P RICHARD , H. M., A LABASTER , T., H ARRIS , N. B. W. & N EARY , C. R. Magmatic Processes Plate Tectonics Geological Society, London, Special Publications, 6, 71– 81. M IDDLEMOST , E. A. K. 1994. Naming material in the magma/igneous rock system. Earth Science Reviews, 37, 215 –224. P AN , G. Q., L U , X. C. & Y U , H. B. 2001. Petrological and geochemical characteristics of Mesozoic adakite from northern Huaiyang and discussion on its genesis. Acta Petrologica Sinica, 17, 541– 500 [in Chinese with English abstract]. Q IAN , Q., C HUNG , S. L., L EE , T. Y. & W EN , D. J. 2002. Geochemical characteristics and petrogenesis of the Badaling high Sr –Ba granitoids: a comparison of igneous rocks from North China and the Dabie–Sulu orogen. Acta Petrologica Sinica, 18, 275 –292.

Q IU , J. S., W ANG , D. Z., Z ENG , J. H. & M CE I NNES , B. I. A. 1997. Study on trace element and Nd–Sr isotopic geochemistry of the Mesozoic K-rich volcanic rocks and lamprophyres in western Shandong. Geological Journal of China Universities, 3, 384– 395 [in Chinese]. R APP , R. P. & W ATSON , E. B. 1995. Dehydration melting of metabasalt at 8 –32 kbar: implications for continental growth and crust–mantle recycling. Journal of Petrology, 36, 891– 931. R APP , R. P., S HIMIZU , N., N ORMAN , M. D. & A PPLEGATE , G. S. 1999. Reaction between slab-derived melts and peridotite in the mantle wedge: experimental constraints at 3.8 GPa. Chemical Geology, 160, 335–356. R OGERS , N. W., J AMES , D., K ELLEY , S. P. & DE M ULDER , M. 1998. The generation of potassic lavas from the eastern Virunga province, Rwanda. Journal of Petrology, 39, 1223–1247. R OLLINSON , H. 1993. Using Geochemical Data: Evaluation, Presentation, Interpretation. Longman, Harlow, 108– 111. R OWLEY , D. B., X UE , R., T UCKER , R. D., P ENG , Z. X., B AKER , J. & D AVIS , A. 1997. Ages of ultrahighpressure metamorphism and protolith orthogneisses from the eastern Dabieshan: U–Pb zircon geochronology. Earth and Planetary Science Letters, 151, 191–203. R USHMER , T. 1991. Partial melting of two amphibolites: contrasting experimental results under fluid-absent conditions. Contributions to Mineralogy and Petrology, 107, 41– 59. S EN , C. & D UNN , T. 1994. Dehydration melting of a basaltic composition amphibolite at 1.5 and 2.0 GPa: implications for the origin of adakites. Contributions to Mineralogy Petrology, 117, 394– 409. S ENGO¨ R , A. M. C. & N ATAL ’ IN , B. A. 1996. Paleotectonics of Asia: fragments of a synthesis, In: Y IN , A. & H ARRISON , T. M. (eds) The Tectonic Evolution of Asia. Cambridge University Press, Cambridge, 486–641. S UN , J. G., L IU , C. H. & Z HENG , C. Q. 1996. Genetic mechanism of the Mesozoic intermediate– acid granitic complex in eastern Jiaodong area. Geological Journal of China Universities, 2, 207–217. W ANG , L. G., Q IU , Y. M., M C N AUGHTON , N. J., G ROVES , D. I., L UO , Z. K., M IAO , L. & L IU , Y. K. 1998. Constraints on crustal evolution and gold metallogeny in the northwestern Jiaodong Peninsula, China, from SHRIMP U– Pb zircon studies of granitoids. Ore Geology Review, 13, 243– 258. W ANG , Q., X U , J. F. & Z HAO , Z. H. 2001. The petrogenesis and geodynamic significance of HREE depleted granitoids during Yanshan period in the Dabie Mountains. Acta Petrologica Sinica, 17, 551–564 [in Chinese with English abstract]. W OLF , M. B. & W YLLIE , J. P. 1994. Dehydration-melting of amphibolite at 10 kbar: the effects of temperature and time. Contributions to Mineralogy and Petrology, 115, 369– 383. X U , J. F., S HINJO , R., D EFANT , M. J., W ANG , Q. & R APP , R. P. 2002. Origin of Mesozoic adakitic intrusive rocks in the Ningzhen area of east China: partial

Polyphase Mesozoic tectonics in the eastern part of the North China Block: insights from the eastern Liaoning Peninsula massif (NE China) W. LIN1, M. FAURE2, P. MONIE´3 & Q.-C. WANG1 1

State Key Laboratory of Lithospheric Evolution, Institute of Geology and Geophysics, Chinese Academy of Sciences, Beijing 100029, China (e-mail: [email protected]) 2

Institut des Sciences de la Terre d’Orle´ans, UMR – CNRS 6113, Baˆtiment Ge´osciences, Universite´ d’Orle´ans, F-45067 Orle´ans Cedex 2, France 3

Laboratoire de Dynamique de la Lithosphe`re, UMR – CNRS 5573, Universite´ de Montpellier 2 Pl. E.-Bataillon, F-34095 Montpellier Cedex, France Abstract: In Eastern China, the North China Block (NCB) has experienced a complex tectonic evolution during Late Palaeozoic to Mesozoic times. The unconformable sedimentary cover, which ranges in age from Neoproterozoic to Permian, underwent two tectonic episodes in Early Triassic and Cretaceous times. An early north–south compressional phase, D1, characterized by north-verging recumbent folds with east–west-trending axes and top-to-the-north ductile shear zones can be observed in a gabbroic pluton and the Neoproterozoic sedimentary host rocks. The available radiometric ages allow us to interpret this event as an Early Triassic back-thrusting related to the final stage of the collision between the North China and South China Blocks. During the Late Mesozoic, an extensional event characterized by (1) half-grabens filled by continental terrigeneous red beds and lava flows, (2) low-angle detachment faults, (3) synkinematic granitic plutons, and (4) metamorphic core complexes is widespread in the Eastern Liaoning Peninsula. This extensional D2 event can be subdivided into a ductile deformation and a brittle deformation, corresponding to different expressions of crustal deformation. In every area studied in the Eastern Liaoning Peninsula, the ductile D2 deformation is characterized by a NW–SE-trending stretching lineation with a top-to-the-NW sense of shear. Mica and amphibole from the metamorphic country rocks, granodioritic or monzogranitic plutons and their mylonitized margins yield 40Ar/39Ar ages ranging from 130 to 120 Ma. These dates support fast cooling and exhumation rates coeval with the extensional tectonics. The brittle D2 deformation is responsible for the formation of high-angle normal faults marked by low-temperature cataclasites bounding half-grabens. During the Mesozoic, the tectonic regime of Eastern China experienced a significant inversion from compression to extension. The Eastern Liaoning Peninsula massif provides an example of this geodynamic transition.

Along the eastern margin of Eurasia, the North China Block (NCB) experienced several tectonic events during the Phanerozoic. The south margin of the NCB, the Qinling–Dabie–Sulu Mountain Belt, corresponds to the region of the collision between the NCB and the South China Block (SCB). However, the timing of this collision remains disputed, with dates ranging from Early Palaeozoic to Late Triassic (e.g. Mattauer et al. 1985, 1991; Okay et al. 1993; Hacker et al. 1995; Meng & Zhang 1999; Wallis et al. 2005). The widespread presence of inclusions of coesite and microdiamond in continental crustal rocks indicates important continental subduction (Wang et al. 1989; Yang & Smith 1989; Xu et al. 1992). According to Fan & Menzies (1992), the plate convergence was accompanied by more than 200 km of thickening of the East China lithosphere. Along the north margin of the NCB, the geometries of the multiple collisions of microplates or island

arcs between the NCB and Mongolia that occurred during Late Permian to Early Triassic times are not yet settled (e.g. Wang & Liu 1986; Yin & Nie 1996; Xiao et al. 2003). Eclogites indicating a syncollision high-pressure metamorphism was recently discovered (Ni et al. 2004). During the Late Mesozoic, widespread continental extension characterized by volumetrically important Jurassic–Cretaceous plutonism and volcanism, extensional basins and several metamorphic complexes has long been recognized (Huang 1945). During this period, the lithospheric keel lost more than 120 km (Menzies et al. 1993). However, the geodynamic significance of this Mesozoic extensional deformation is still poorly understood and is disputed (Yin & Nie 1993; Allen et al. 1997; Ratschbacher et al. 2000; Davis et al. 2001; Ren et al. 2002; Meng et al. 2003). Our work in the Eastern Liaoning Peninsula

From: ZHAI , M.-G., WINDLEY , B. F., KUSKY , T. M. & MENG , Q. R. (eds) Mesozoic Sub-Continental Lithospheric Thinning Under Eastern Asia. Geological Society, London, Special Publications, 280, 153–170. DOI: 10.1144/SP280.7 0305-8719/07/$15 # The Geological Society of London 2007.

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massif focuses on structural and geochronological aspects of the Cretaceous extension. Moreover, evidence for a Triassic compression is also provided. The compressional and extensional events are discussed in the frame of the collision of the NCB and SCB and their subsequent tectonic evolution.

Geological framework of the Eastern Liaoning Peninsula In the east of Liaoning Province of Northeastern China, the Eastern Liaoning Peninsula massif consists of a large metamorphic and granitic complex. Early Precambrian foliated granite, gneiss, migmatite and metasedimentary rocks occupy almost half

of the massif (Fig. 1). Neoproterozoic and Palaeozoic sedimentary rocks are distributed along the margin of the metamorphic and magmatic complex. Numerous Palaeoproterozoic and Mesozoic granitoid plutons intrude the massif (Wu et al. 2005a, b). Late Mesozoic to Cenozoic tuffaceous and terrestrial red sandstone and conglomerate fill several half-grabens and pull-apart basins formed in response to regional extensional tectonics (Allen et al. 1997; Ren et al. 2002). The geological framework of the Eastern Liaoning Peninsula massif can be divided into two. In the north, a Palaeoproterozoic orogen, composed of mafic –ultramafic rocks and marine sedimentary rocks, crops out between a northern Archaean

Fig. 1. Generalized geological map of Eastern Liaoning Peninsula. Inset shows the location of the study area in East Asia. NCB, North China Block; SCB, South China Block; TLF, TanLu Fault; extensional domes formed around 110–145 Ma: Y, Yiwulushan dome (Ma et al. 1999; Darby et al. 2004); YM, Yunmengshan dome (Davis et al. 1996); H, Hohhot dome (Davis et al. 2002; Wang et al. 2004); D, Dabieshan dome (Hacker et al. 1998; Faure et al. 1999); W, Wugongshan dome (Faure et al. 1996), L, Lushan dome (Lin et al. 2000).

POLYPHASE MESOZOIC TECTONICS

block and a southern Palaeoproterozoic block (Faure et al. 2004). The mafic magmatic belt is interpreted as an arc developed above a southdirected subduction zone, which was subsequently thrust to the north upon the Archean block (Fig. 1). In the south, several antiformal metamorphic areas occur, the largest of which is a Late Mesozoic metamorphic core complex that crops out NE of Jinzhou (Liu et al. 2005; Lin et al. 2006). In this paper, we provide results of field work and geochronological constraints obtained by the 40 Ar/39Ar method from several sites of the East Liaoning Peninsula (Fig. 1). These data show that the region experienced ductile and brittle deformations related to the Cretaceous extensional tectonics. Also, evidence for an earlier compression can be recognized in the southern part of the peninsula.

Cretaceous extensional tectonics in the Eastern Liaoning Peninsula In this section, several representative examples of brittle and ductile structures formed during the

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Cretaceous in the northern, central and southern parts of the Eastern Liaoning Peninsula are described.

Extensional structures in the Xiuyan area In the central part of the Eastern Liaoning Peninsula, several well-developed extensional structures occur in the Xiuyan area (Fig. 2). Rare plant and bivalve fossils indicate that a weakly metamorphosed sandstone, conglomerate, phyllite and slate unit was deposited in the Permian (LBGMR 1989). A mylonitic zone separates this weakly metamorphosed Permian series from the underlying metamorphic rocks (Fig. 3). An N120–150E-trending mineral and stretching lineation marked by biotite or muscovite streaks, aligned andalusite crystals, and elongated or boudinaged clasts (Fig. 4) displays evidence of top-to-the-NW shear along the contact. The Permian rocks tectonically overlie Proterozoic metamorphic rocks. A muscovite–biotite schist (LN141), sampled in this mylonitic zone, is dated from this locality (see below). To the SW of this Permian massif, a Cretaceous porphyritic granodiorite and an adjacent Late

Fig. 2. Geological map of Xiuyan and Buyunshan area (see Fig. 1 for location) with emphasis on brittle and ductile extensional structures discussed in the text. LN134, LN135 and LN141 are the samples dated by 40Ar/39Ar analysis (see Fig. 5).

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Fig. 3. Cross-section through a Cretaceous half-graben, granite and Permian unit east of Xiuyan (location is shown in Fig. 2).

POLYPHASE MESOZOIC TECTONICS

Jurassic to Early Cretaceous basin are also involved in the regional extension (Fig. 2; LBGMR 1989). The western border of this porphyritic granodiorite develops a strong top-to-the-NW shear associated with a low-temperature mylonitic fabric over a few tens of metres (Fig. 4). Two samples are dated from here (see below): a granite mylonite (LN134) and an undeformed granite (LN135), which are 200 m apart (Figs 2 and 3). To the west, the ductile structures are overprinted by a high-angle brittle normal fault that forms the eastern margin of the Cretaceous basin. Although a detailed sedimentological study has not been conducted here, the progressive tilt of the bedding and the west-to-east coarsening of the sandstone and conglomerate deposits are consistent with synsedimentary activity of the normal fault. On the fault surface, the down-dip slickenlines are parallel to the stretching lineation and show that these brittle and ductile deformations developed during the same extensional event (Fig. 3). To assess the time of this extensional tectonics, samples LN134, LN135 and LN141 were collected for laser probe 40Ar/39Ar dating using a single-grain step-heating procedure. Two biotites from the extensional shear zone and the granodiorite (samples LN134 and LN135) yield very similar plateau ages of 122.2 + 1.2 Ma and 122.8 + 1.2 Ma, respectively (Fig. 5) and identical intercept ages on an isotope correlation plot. A minor amount of excess argon is released in the first heating increments, representing a loosely bound component trapped at the grain surface or along mineral defects. In the basal shear zone of the Permian schist (Fig. 3), biotite and muscovite were also analysed. Biotite from sample LN141 has an age spectrum similar to that observed in the two former biotites of LN134 and LN135 (Fig. 5). After release of excess argon at low experimental temperature, consistent ages at 129.9 + 1.3 Ma are obtained for 90% of the argon released. A similar intercept age is given by the isotope correlation plot, with an atmospheric composition for initial argon. Muscovite from the same sample has a flat age spectrum (Fig. 5) with a plateau date of 127.6 + 1.4 Ma and an intercept age of 128.9+1.3 Ma in the isotope correlation plot, with a low initial argon ratio suggesting an episode of argon loss. These 122–130 Ma 40Ar/39Ar biotite and muscovite ages allow us to conclude that the ductile normal faulting related to the regional extensional event took place during the Early Cretaceous.

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Extensional structure in the Sinian Series of the Buyunshan area SW of Xiuyan, another low-angle detachment fault is exposed in the Buyunshan area, where it separates Neoproterozoic sedimentary and Palaeoproterozoic metamorphic rocks in the hanging wall and footwall, respectively (Fig. 2). Along the contact, ductile –brittle to brittle structures are conspicuous. The lithological variety was responsible for distinct rheological behaviours (Fig. 6); pelites were ductilely sheared whereas sandstones were fractured. A flat-lying slaty cleavage with a NW–SE stretching lineation is marked by quartz fibres, and chlorite –sericite clots or quartz pressure shadows developed around pyrite. In massive sandstone layers, en echelon tension gashes and Riedelshears form consistently with the ductile–brittle shearing (Fig. 6). In some places, the slaty cleavage is deformed with a sigmoidal shape and in others centimetre-scale sandstone– mudstone alternations are folded with axes nearly perpendicular to the shearing (Fig. 2). Because of the flat-lying attitude of shear surfaces, the extensional or compressional character of these structures is difficult to assess. However, as the contact superimposes younger (i.e. Neoproterozoic rocks) above older (Palaeoproterozoic) rocks, an extensional setting is preferred here. The top-to-the-NW displacement agrees well with other extension-related structures observed in the Eastern Liaoning Peninsula.

Extensional structures in the Liaoyang – Benxi area At the northern end of the Eastern Liaoning Peninsula, extensional shearing is recognized south of Liaoyang and Benxi (Figs 1 and 7). South of Benxi, flat-lying Neoproterozoic sandstone unconformably covers Palaeoproterozoic rocks that constitute a Palaeoproterozoic belt (Faure et al. 2004, and references therein). North of an ENE – WSW-trending normal fault, the Neoproterozoic – Palaeozoic rocks are folded and faulted. South of Liaoyang, the ENE –WSW normal fault connects with a NW –SE strike-slip fault. Thus, the bulk geometry of the Neoproterozoic– Palaeozoic series appears as an extensional allochthon in which second-order normal faults develop. The sedimentary rocks are deformed by NE –SW-trending

Fig. 4. Field-scale photographs showing extensional Cretaceous deformation (location of photographs is shown in Fig. 2). (a) Highly sheared Permian schists, east of Xiuyan. (b) Stretching lineation of metapelite (phyllite) in a basal shear zone between Permian rocks and Precambrian basement, east of Xiuyan, corresponding to sample LN141 dated by 40Ar/39Ar analysis (see Fig. 5). (c) Cretaceous porphyritic granodiorite with K-feldspar porphyroclast pressure shadow and shear band. (d) S –C mylonite with a top-to-the-NW sense of shear associated with a normal ductile fault in a Cretaceous granodiorite. (e) Layer slip on bedding in Sinian sandstone.

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Fig. 5. 40Ar/39Ar age spectra of biotite (a –c) and muscovite (d), from mylonitic and undeformed granite, and pelitic schist from the Xiuyan area (see Figs 2 and 3 for location).

folds overturned to the NW (Fig. 8). The subhorizontal or low-angle axial planes agree well with a folding developed coevally with normal faulting. Likewise in the southern areas of the Eastern Liaoning Peninsula, the age of normal faulting is inferred from a Cretaceous half-graben that crops out east of Benxi (Fig. 7).

Extensional structures in the south of the Eastern Liaoning Peninsula massif In the southern part of the Eastern Liaoning Peninsula, the bulk architecture of the massif is dominated by an ENE –WSW-trending dome that records a synmetamorphic deformation, defined here as D2 (Fig. 9a). Our detailed work allows us to interpret this massif as a Mesozoic metamorphic

Fig. 6. Examples of Cretaceous extensional deformation in the Neoproterozoic (Sinian) series in the Buyunshan area (redrawn from pictures) (location of figures is shown in Fig. 2). (a) Heterogeneously sheared sandstone–mudstone alternation; (b) sigmoidal cleavage in pelite; (c) folded sandstone with sheared limbs.

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Fig. 7. Geological map and cross-section of the Liaoyang–Benxi area. South of Liaoyang– Benxi, the Neoproterozoic–Palaeozoic series is deformed by north-verging drag folds related to an extensional allochthon (see Fig. 1 for location).

core complex formed in response to NW– SE extension (Lin et al. 2006). NE of Jinzhou city, Proterozoic gneiss and foliated migmatite that form the core of the dome are heterogeneously deformed with a relatively weakly foliated core and mylonitic shear zones at the margin (Fig. 9b). The dome boundary consists of a ductile detachment normal fault (Fig. 9b). Early Cretaceous syntectonic plutons that are involved into this detachment fault are converted to mylonite or ultramylonite. This detachment fault strikes NNE –SSW and dips northwestward in the western part, and strikes ENE– WSW and dips to ESE in the eastern part. East of Jinzhou city, the Neoproterozoic sedimentary rocks crop out in the hanging wall of the detachment fault (Fig. 9a). As observed in the typical metamorphic core complexes of North America (Lister & Davis 1989), the detachment fault of the East Liaoning Peninsula massif is arched as a result of synextensional folding with a NE–SW axis (Fig. 9b). As a result, the SE dome limb appears as an apparent top-to-the-NW thrust, but is in reality a folded low-angle normal fault. In the hanging wall

of the detachment fault, the Neoproterozoic and Palaeozoic sedimentary rocks are deformed by northwestward-verging folds (Fig. 9b). The planar structures exhibit a conspicuous mineral and stretching lineation with a dominantly NW –SE trend (Fig. 9c). Along the detachment normal fault, mylonites are commonly developed in the Mesozoic granite and migmatitic basement. Near the dome margin, the migmatite and Mesozoic granite are reworked into mylonite or even ultramylonite (Fig. 9b). A top-to-the-NW shearing is unambiguously shown by quartz, biotite and amphibole pressure shadows, feldspar or amphibole sigmaor delta-type porphyroclast systems, and shear bands (Fig. 10b). In tonalitic granite, S–C structure matches other kinematic criteria. The leucosome and mafic restites of the migmatite are sheared with sigmoidal shapes (Fig. 10a). The leucosomes is also asymmetrically folded or boudinaged by this top-to-the-NW shearing. In the south and SE parts of this metamorphic core complex, metamorphism is absent or very weak in the Neoproterozoic and Palaeozoic series, but a slight

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Fig. 8. Field-scale photographs showing Cretaceous extensional deformation in the Liaoyang–Benxi and Buyunshan areas (location of photographs is shown in Figs 2 and 7). (a) Decametre-scale folds overturned to the north in Neoproterozoic turbidite. (b) Foliated cataclasite in Palaeoproterozoic micaschist. Sigmoidal quartz veins and lenses indicate top-to-the-north shear. (c) Layer slip and oblique fracture in Palaeozoic limestone. (d) Decametre- to centimetre-scale folds overturned to the north in Palaeoproterozoic micaschist overlain by the Liaoyang–Benxi extensional allochthon. (e) Subhorizontal brittle–ductile D2 shear in folded Neoproteorozoic schist in the Buyunshan area.

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Fig. 9. Structure of the Jinzhou area. (a) Structural map of the Jinzhou dome (location shown in Fig. 1). North of Jinzhou, the east–west-trending folds and thrust related to a pre-extension top-to-the-north compression should be noted (see text for discussion). (b) Cross-section drawn parallel to the direction of the main stretching lineation (the pluton roots are hypothetical) showing ductile and brittle extensional structures. (c) Stereoplot of foliation and lineation.

recrystallization of sericite or chlorite may occur in the pelitic layers. However, in some parts of these sedimentary series, bedding is overprinted by a slaty cleavage and a NW–SE direction of mineral

and stretching lineation marked by elongated and recrystallized chlorite and quartz grains (Fig. 10c). Along this lineation, top-to-the-NW shearing is clearly indicated by sericite and chlorite shear

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Fig. 10. Photographs of extensional structures of Late Mesozoic age (location of photographs is shown in Fig. 9). (a) Sigmoidal leucosome and shear bands in migmatite showing top-to-the-NW shear. (b) Amphibole around feldspar clast in the mylonitic margin of a Cretaceous granite showing top-to-the-NW shear. (c) Brittle shear in Neoproterozoic mudstone of the dome cover. The sigmoidal shape of the lenses indicates top-to-the-NW shear. (d) Chlorite developed in extensional shear bands in weakly metamorphosed Neoproterozoic sedimentary cover with top-to-the-NW shear.

bands in thin section (Fig. 10d). Thus, the deformation responsible for the exhumation of the South Liaodong Peninsula massif is well recorded both in the sedimentary cover and in the metamorphic core. Along the ductile detachment normal fault, several samples were collected to determine the age of exhumation of the metamorphic core complex in the south of the Eastern Liaoning Peninsula. A single-grain step-heating procedure was carried out by the laser probe 40Ar/39Ar method. Four samples were taken on both the northwestern and southeastern sides of the dome to place age constraints on the top-to-the-NW shearing that is the prominent deformation feature. On the northwestern side of the dome, amphibole from a mylonitic Late Jurassic granodiorite (Wu et al. 2005b) and biotite from the gneissic basement give 40Ar/39Ar plateau ages ranging from 121.7 + 1.6 Ma to 116.4 + 1.5 Ma (Lin et al. 2006). On the southeastern side of the dome, biotites from mylonitic Cretaceous monzogranite and migmatite yield plateau ages from 117.7 + 1.2 Ma to 116.9 + 1.4 Ma. These 116 –122 Ma ages for the detachment fault indicate an Early Cretaceous age for this ductile extensional event.

SE of Wafangdian City, a west-dipping highangle brittle normal fault overprints the ductile fault (Fig. 9), and controls the opening and sedimentary infill of a Mesozoic half-graben. Sedimentary rocks in the basin yield Early Cretaceous fossils (LBGMR 1989). This basin is similar to other half-graben basins found throughout the Eastern Liaoning Peninsula, such as those in the Xiuyan area (see above) or in the Zhuanghe and Dandong areas (Fig. 1). Extensional basins overlie the mylonitic rocks formed at an early stage during the exhumation of the metamorphic core complex. Although both ductile and brittle faulting events are due to the same tectonic phenomena, our structural observations allow us to derive a relative timing. The brittle extensional deformation clearly overprints the ductile one. In conclusion, in the Eastern Liaoning Peninsula, brittle normal faults with down-dip slickenlines are widespread. Intensely sheared decametre-scale cataclasite and breccia zones indicate that the Late Mesozoic extension is regionally significant. Late Mesozoic basins and granitoids formed in an extensional setting can be observed not only along the boundary but also in the central part of the massif.

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All of these Late Mesozoic continental basins are half-grabens bounded by high-angle brittle normal faults that control their opening and sedimentary infill. The largest basins crop out in areas east of Wafangdian, at Zhuanghe, at Dandong, east of Xiuyan and at Benxi (Fig. 1). On the other margin of each basin, continental clastic red beds unconformably overlie the metamorphic rocks and Neoproterozoic– Palaeozoic series. Inside the basins, andesitic or basaltic lava flows and interstratified tuff layers suggest that magma channelling took place during basin opening. High-angle brittle normal faults are commonly superimposed on mylonitic zones developed at the expense of Precambrian basement rocks or Mesozoic granitoids. Two generations of Mesozoic plutons are recognized in the Eastern Liaoning Peninsula. The oldest and youngest are dated as Early–Middle Jurassic and Early Cretaceous, respectively (Wu et al. 2005a, b). The Cretaceous plutons exhibit conspicuous planar and linear mineral preferred orientations in their margins, but almost isotropic textures in their cores. Microscopic investigations show that the granite fabric was formed under subsolidus conditions; that is, when the magma started to crystallize. At the pluton margin, there is post-solidus hightemperature mylonite of a few tens metres thickness, characterized for instance by dynamic recrystallization of quartz grains. The progressive evolution from unfoliated granite to well-foliated mylonitic margins shows that the plutons are synkinematic bodies, which acquired their internal structure during crystallization. Thus the granite fabrics, and particularly those at the pluton margin, provide geometric and kinematic information on the deforming host rocks at the time of granite emplacement. In all these studied plutons, the mineral and stretching lineation trends NW–SE (Fig. 1). In a section parallel to the lineation and perpendicular to the foliation, kinematic criteria such as sigmoidal biotite, sigma- or delta-type porphyroclast systems or S–C structures consistently indicate a top-to-the-NW motion of the country rock with respect to the pluton. This kinematic pattern indicates an extensional setting at the time of pluton emplacement. The older pluton generation consists of Jurassic granodiorite or tonalite. Conversely to the Cretaceous ones, the Jurassic plutons were deformed under postsolidus conditions. Therefore, from our study at several places in the Eastern Liaoning Peninsula, two different crustal expressions of the Cretaceous extensional deformation are recognized. The brittle event formed in the upper crust is related to the formation of Cretaceous half-basins. The ductile event is coeval with the emplacement of Cretaceous monzogranitic plutons and exhumation of the Palaeoproterozoic to Archaean basement. Because of their

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similar geometric and kinematic patterns and age, these ductile and brittle deformations belong to a single phase of Cretaceous extensional tectonics.

Evidence for an Early Mesozoic compression In the southernmost part of the Eastern Liaoning Peninsula, several structural observations argue for compressional tectonics (Figs 1 and 10). The most convincing evidence is provided by a gabbroic pluton intruding Neoproterozoic sandstone. The gabbroic pluton is also intruded by a syenite pluton in which zircons yields U –Pb ages of c. 219 Ma (Wu & Yang 2005; figs 11 and 12a). The margins of the gabbroic pluton and its Neoproterozoic host rocks are pervasively foliated and lineated. The south to SW dip of the foliation shows that the gabbro lies in the hanging wall of a thrust fault (Fig. 11b). Along the north –south- to NE –SW-trending mineral and stretching lineation, top-to-the-north sigma-type feldspar, biotite and amphibole porphyroclasts or quartz pressure shadows are observed both at field and thin-section scales (Fig. 12b and c). A few hundred metres north of the contact, north-verging folds associated with an axial planar cleavage deform the Neoproterozoic series. Another thrust zone associated with folds overturned to the north is recognized north of Jinzhou City (Figs 1 and 9). Along the western margin of the metamorphic core complex described in the previous section, nearly east –west-trending northverging recumbent folds and north-directed brittle thrust faults argue for a north–south shortening. The south-dipping axial planar cleavage developed in the Neoproterozoic rocks, and in the Palaeozoic limestone and sandstone, is in agreement with this north-directed horizontal shearing (Fig. 12d–f ). It is worth noting that, east –west-trending folds are cut by the above-described ductile detachment extensional shear zone. All these observations indicate that the Neoproterozoic to Late Palaeozoic sedimentary cover of the southern part of the Eastern Liaoning Peninsula massif experienced a bulk north–south shortening and northward thrusting before the Cretaceous extensional tectonics. Because of the lack of syntectonic metamorphism, it is difficult to directly determine the time of this compressional deformation. The youngest deformed rocks, which crop out north of Jinzhou City, are Permian limestone (LBGMR 1989). The continental Cretaceous red beds and exhumed metamorphic complex overprint this north–south folding (Fig. 9). Thus the age of this deformation should be bracketed between Late Palaeozoic and Early Mesozoic. Moreover, zircons of the

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undeformed syenite pluton that cuts the north thrust fault yield a U –Pb age of 219 + 1 Ma (Wu & Yang 2005). This indicates that the compressional event occurred after the Permian and before the Late Triassic. These lines of evidence allow us to infer an Early–Middle Triassic age.

Discussion The Eastern Liaoning Peninsula massif is a structurally complex area that experienced several superimposed deformation events. When dealing with the Mesozoic tectonics, the earliest event, D1, of

Early–Middle Triassic age, corresponds to top-to-the-north (or NE) shearing coeval with crustal thickening (Yin & Nie 1993, 1996). In the south of the massif, recumbent folds involving the Neoproterozoic –Palaeozoic sedimentary sequence and ductilely deformed gabbro associated with a north-directed thrust over the Neoproterozoic sedimentary rocks are also attributed to this D1 compression. The ductile D2 top-to-the-NW shearing is associated with the exhumation of the metamorphic basement and the intrusion of Mesozoic plutons. This event is mainly observed in the metamorphic core complex of the Eastern Liaoning Peninsula

Fig. 11. Structural map of the southernmost part of the Eastern Liaoning Peninsula (a) and a cross-section (b) drawn parallel to the direction of the stretching lineation. An Early–Middle Triassic gabbro overthrusts the Neoproterozoic sedimentary series. This structure is interpreted to relate to an early stage of compression (the pluton roots are hypothetical).

Fig. 12. Evidence for early compressional deformation (location of photographs is shown in Figs 9 and 11). (a) Undeformed Late Triassic syenite intruding a foliated and lineated gabbro. (b) Sigmoidal feldspar (F) and amphibole (A) in gabbro mylonite showing top-to-the-N shear. (c) Amphibole and biotite pressure shadow around plagioclase clast in gabbro showing top-to-the-NE shear. (d) Field photograph showing north-vergent recumbent folds and thrust faults in Neoproterozoic sandstone, NW of Jinzhou city (location shown in Fig. 8). (e, f) Detail of axial planar cleavage at metre and centimetre scales, respectively.

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Massif. The Buyunshan flat-lying fault and the extensional allochthon south of Liaoyang–Benxi also belong to D2. The brittle D2 Cretaceous extensional event is responsible for the formation of continental half-grabens. These deformations leading to the final bulk geometry of the Eastern Liaoning Peninsula massif are summarized in Figure 13.

Back-thrust interpretation of the Triassic compression The regional tectonic significance of this northward thrusting and folding event remains hypothetical. According to Yin & Nie (1996, and references therein), top-to-the-north shearing might be related to the indentation of the NCB by the SCB during Triassic collision. During the Early Triassic, a compressional regime related to the northward subduction of the South China Block below the North China Block is well acknowledged (e.g. Mattauer et al. 1985, 1991; Okay et al. 1993; Hacker et al. 1995). Simultaneously, northward thrusting also occurred within the continental

crust of the overriding North China Block (Hacker et al. 1998; Wang et al. 1998; Faure et al. 2001, 2003). This large-scale deformation is not restricted to the Qinling –Dabie –Sulu belt, but occurs throughout the southern margin of the North China Block. Thus, in the Eastern Liaoning Peninsula massif, we interpret the Triassic compressional tectonics as back-thrust related to the final collision between the NCB and the SCB.

Regional significance of the Mesozoic extensional deformation in Eastern China During the Mesozoic, the tectonic regime of Eastern China experienced a drastic change, as it shifted from a Late Permian–Early Triassic compression to a Jurassic–Cretaceous extension (Zhao et al. 1994, 2004; Zhai et al. 2004). The Cretaceous features recognized in the Eastern Liaoning Peninsula agree well with those described in many places in Eastern China, where Jurassic–Cretaceous continental basins filled by sandstone, mudstone and conglomerate interbedded with acidic or mafic

Fig. 13. Synthetic block diagram of the Southern part of the Eastern Liaoning Peninsula massif showing the bulk geometry and polyphase deformation. Top-to-the-north D1 Triassic compression is followed by ductile and brittle Cretaceous extension.

Fig. 12. Evidence for early compressional deformation (location of photographs is shown in Figs 9 and 11). (a) Undeformed Late Triassic syenite intruding a foliated and lineated gabbro. (b) Sigmoidal feldspar (F) and amphibole (A) in gabbro mylonite showing top-to-the-N shear. (c) Amphibole and biotite pressure shadow around plagioclase clast in gabbro showing top-to-the-NE shear. (d) Field photograph showing north-vergent recumbent folds and thrust faults in Neoproterozoic sandstone, NW of Jinzhou city (location shown in Fig. 8). (e, f) Detail of axial planar cleavage at metre and centimetre scales, respectively.

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Mesozoic tectonics in the Eastern Block of the North China Craton: implications for subduction of the Pacific plate beneath the Eurasian plate S. Z. LI1, T. M. KUSKY1,2, G. ZHAO3, F. WU4, J.-Z. LIU5, M. SUN3 & L. WANG1 1

Department of Marine Geology, Ocean University of China, Qingdao 266100, China (e-mail: [email protected]) 2

Department of Earth and Atmospheric Sciences, St Louis University, 3507 Laclede Avenue, St Louis, MO 63103, USA

3

Department of Earth Sciences, The University of Hong Kong, Hong Kong, China

4

Institute of Geology and Geophysics, Chinese Academy of Sciences, Beijing 100029, China

5

National Astronomical Observatories, Chinese Academy of Sciences, Beijing 100012, China Abstract: The Jiao-Liao massif is located in the hanging wall of the north-dipping Dabie–Sulu suture zone and is an important part of the Eastern Block of the North China Craton. Several important tectonic models for the tectonic evolution of Eastern Asia rely on critical information from the Jiao-Liao massif. This paper combines new sensitive high-resolution ion microprobe (SHRIMP) U–Pb zircon ages of the Dandong Granite in the southern Liaoning Province, China, with extensive field data for the eastern North China Craton, including the Bohai Bay Basin. Combined with other recent SHRIMP dating, we use this information to summarize the Mesozoic tectonic reactivation and evolutionary processes of the Jiao-Liao massif of the Eastern Block of the North China Craton. In this study we identify a c. 160 Ma episode of partial melting of Palaeoproterozoic plutons in the Jiao-Liao massif. Cathode luminescence and backscatter electron imagery reveal c. 167–157 Ma magmatic euhedral single zircons and magmatic zircon rims surrounding c. 2100 Ma cores in the Dandong Granites near the Liaonan Neoarchaean terrane. This partial melting is probably related to in situ remelting of ancient lower continental material, mostly the North China Craton. The Dandong plutons are aligned in a NE–SW direction and are extensively deformed by subhorizontal ductile thrust-related shearing and subsequent NNE– SSW trending folds. Here, we show that for the Dandong area the first deformation occurred between 195 and 193 Ma, based on K–Ar and 40Ar/39Ar ages of muscovites from east– west-trending shear zones on the Liaodong Peninsula. Based on the field relationships between the plutons and structural fabrics, a range from 153 to 145 Ma is defined as the duration of the second deformation in the Dandong Granites. The third deformation is marked by the formation of NNE– SSW strike-slip faults between 135 and 95 Ma. This deduced age range is similar to an 40Ar/39Ar age range of 128–132 Ma of initial sinistral strike-slip faulting of the Tan-Lu fault in Anhui Province and to a biotite cooling age of 100 + 2.3 Ma of the Yilan– Yitong segment of the Tan-Lu fault in the Jilin Province. These faults are transtensive and controlled the formation of pull-apart basins. However, during the third deformation, some metamorphic core complexes in Eastern China formed in the overlapping area between the largescale sinistral faults. Our SHRIMP data also indicate that the Liaodong basement and its Early Mesozoic magmatism are similar to the Jiaodong basement and its Mesozoic magmatism. Therefore, the Early Mesozoic evolution of the Liaodong area, similar to that of the Jiaodong area, was also closely related to the Sulu orogen in the Early Mesozoic and to the Pacific subduction throughout the Mesozoic.

The North China Craton has had a long tectonic history with major tectonothermal events in the early Archaean, late Archaean, NeoProterozoic, Mesozoic, and Cenozoic (Zhao et al. 1999, 2000, 2001, 2002; Zhao 2001; Kusky & Li 2003; Zhai & Liu 2003; see review by Kusky et al. 2007). The Jiao-Liao massif (Fig. 1) is an important part of the Eastern Block of the North China Craton

located in the hanging wall of the north-dipping Dabie –Sulu suture zone (Li 1994; Yang et al. 2002). Many researchers have attempted to determine to what extent the stable Precambrian block has experienced intensive reworking during collision of the North and South China Cratons and formation of the Dabie –Sulu orogen. Although several important tectonic models for the tectonic

From: ZHAI , M.-G., WINDLEY , B. F., KUSKY , T. M. & MENG , Q. R. (eds) Mesozoic Sub-Continental Lithospheric Thinning Under Eastern Asia. Geological Society, London, Special Publications, 280, 171– 188. DOI: 10.1144/SP280.8 0305-8719/07/$15 # The Geological Society of London 2007.

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Fig. 1. Tectonic division of Eastern China showing Mesozoic thrust faults and folds, and metamorphic core complexes. NCC, North China Craton; WB, Western Block; EB, Eastern Block; QL– DB–SL, Qinling– Dabie–Sulu orogen; SCC, South China Craton; TC, Tarim Craton. 1, Yiwulushan metamorphic core complex (Ma et al. 1999, 2000); 2, Yumeng Shan metamorphic core complex (David et al. 1996); 3, Fangshan metamorphic core complex (Song 1996); 4, Taihanshan metamorphic core complex (Niu 1994); 5, Liaonan metamorphic core complex (the site of the study area, Yang et al. 1996; Chen et al. 1999); 6, Queshan metamorphic core complex (Yang et al. 2000); 7, Western Shandong extensional tectonics (Yan 1994); 8, North Orthogneiss Unit of the Dabie orogen (Ratschbacher et al. 2000); 9, Xingzi metamorphic core complex (Yin & Xie 1996); 10, Wugongshan metamorphic core complex (Faure et al. 1996; Shu et al. 2000a). BJ, Beijing; TCh, Tangcheng; SH, Shanghai; PR Pyeongrang; SE, Seoul.

evolution of NE Asia have been proposed in the last decade, some of which are closely related to the Jiao-Liao massif (Yin & Nie 1993, 1996; Faure et al. 2001; Yang et al. 2002), none systematically describes the responses of the stable Precambrian Eastern Block of the North China Craton to Mesozoic tectonic processes. Yin & Nie (1993) suggested an indentation model, and considered that the north boundary of the Sulu orogen is the Wulian– Yantai fault, north of which Triassic thrust faults and folds are extensively developed (typically in

the hanging wall of the Liaonan detachment fault). However, Faure et al. (2001) proposed a thin-skinned extensional model in which they moved the north boundary of the Sulu orogen to the Bohai Strait or farther to the north, and in which the Jiaobei massif (i.e. the southern Jiao-Liao massif), is equivalent to the basement of the South China Craton. In contrast, Hong (1989) and Yang et al. (2002) proposed a thin-skinned thrust tectonic model. This debate leads us to discuss the Mesozoic tectonics of the Eastern Block of the North China

MESOZOIC TECTONICS IN EASTERN NORTH CHINA

Craton. The geological responses of tectonics and magmatism in the Liaodong Peninsula to the Early Mesozoic collision between the North China plate and the South China plate, and the subduction of the Pacific plate beneath the Eurasian plate, have to date received very little attention. Thus, we concentrate our discussion on Mesozoic tectonics in the Liaodong Peninsula. Important reasons to choose this research area are: (1) there are widespread Triassic to Cretaceous granitoids and multiphase, well-preserved Mesozoic deformation events including thrusting, extension and strike-slip faulting, such as along the Tan-Lu fault system; (2) this area is a link between the Mesozoic Pacific tectonic domain, the Paleozoic– Mesozoic Central Asia tectonic domain and the Paleozoic –Mesozoic Qinling –Dabie tectonic domain; (3) it is a key site to understand ultrahigh-pressure metamorphic (UHPM) processes and to justify pre-existing models that were proposed by geologists and based mainly on ultra highpressure– high-pressure (UHP–HP) geology and/or paleomagnetic data, geochronology, and isotopic geochemistry; (4) the Hefei Basin covers the hanging wall of the Dabie orogen, whereas the JiaoLiao massif preserves excellent outcrops of the hanging wall of the Sulu orogen and is a prime site to study the response to Mesozoic orogeny. In this paper we present our results of sensitive high-resolution ion microprobe (SHRIMP) U –Pb zircon dating of Mesozoic tectonic processes in the Southern Liaoning Province, China. The structural analysis together with our extensive field data for the eastern North China Craton, our work in the oilfields of the Bohai Bay Basin and other recent SHRIMP dating (Miao et al. 1997, 1998; Wang et al. 1998; Luo et al. 1999, 2001) allow us to establish the Mesozoic tectonic processes of the Eastern Block of the North China Craton.

Regional geological setting The Jiao-Liao massif is located in the northeastern part (Zhang & Yang 1988; Li et al. 1998) of the Eastern Block of the North China Craton (Fig. 1; Zhao et al. 1999, 2002). It is bounded by the Central Asia Tectonic Belt to the north (Fig. 1), and the Sulu Triassic orogen and the Gyeonggi Archaean Massif to the south (Li et al. 2001a, b). It comprises Palaeoproterozoic rocks in the centre, Palaeo- to Neo-Archaean basement in the north (Song et al. 1996) and south, and widely outcropping Mesozoic granitoids and Sinian, Paleozoic and Mesozoic strata (Fig. 1). Paleoproterozoic granitoid rocks of c. 2400– 2100 Ma age are sporadically distributed throughout the northerm ranges of the Liaodong Peninsula,

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north of the study area (Fig. 2); (Li & Yang, 1997; Li et al. 1997; Chen et al. 2001; Lu et al. 2004). The Paleoproterozoic Liaohe Group is mainly composed of meta-volcano-sedimentary rocks, and was inhomogeneously metamorphosed in the (widespread) lower greenschist–upper amphibolite facies and (local) granulite facies (Li et al. 1998). The structural fabric of the Paleoproterozoic terrane composed of those rocks generally strikes east –west. Deformed Sinian to Paleozoic strata mainly occur north and SW of the Paleoproterozoic terrane. Mesozoic granitoids crop out in a NE– SW-striking belt and cross-cut the Paleoproterozoic rock series and the folds in the Sinian to Paleozoic strata. Many NE– SW- to NNE-SSW-trending brittle faults offset the Mesozoic plutons and controlled Early Cretaceous pull-apart basins widely seen in the Liaodong, Korean and Jiaodong Peninsulas. Lower Cretaceous strata infilling the pullapart basins mainly consist of 127– 95 Ma andesite, dacite and rhyolite in the lower section (SBGMR 1991) and a coal series in the upper section. No Triassic plutons are known from the Jiaodong Peninsula. However, many Triassic plutons occur in the Liaodong Peninsula and southern Jilin Province. For example, the 229.8 Ma Xiangmo tonalite pluton in the Tonghua region of Jilin Province (Chen 1994). The Dongqin and Yangmulin, Renao, Liangbin and Dahuanggou plutons in the southern Jilin Province have U–Pb ages of 206.7 Ma and 230.7 Ma, and Rb–Sr ages of 239.9 Ma, 234.7 Ma and 242 Ma, respectively (Sui 1995). The Tongjiapu tonalite in the Xiuyan County of Liaoning Province has a biotite K–Ar age of 215.3 Ma. A large pluton of alkaline granite with U–Pb ages of 237–223 Ma occurs in Saima Town of Liaoning Province (Fig. 2; Chen 1994). In the southern Liaodong Peninsula, the Shuangta pluton has a Rb–Sr age of 224 Ma and part of the Yinmawanshan pluton has a Rb–Sr age of 198 + 19 Ma (Yu et al. 2002). Both plutons were deformed by an intensive and subhorizontal ductile shearing event that did not affect the Zhaofang pluton (Rb–Sr age of 140 Ma) or the Guangminshan pluton (U–Pb age of 139.2 Ma) (Yu et al. 2002). The Xiaoheishan pluton was emplaced at about 180 Ma (Yang et al. 2004). Granites in the Liaonan area are mainly latest Early Cretaceous (Chen 1994), including those at Gudaoling (laser ablation–inductively coupled plasma–mass spectrometry (LA-ICPMS) 122 + 2 Ma), Yinmawanshan (LA-ICPMS 122 + 6 Ma), Qianshan (LA-ICPMS 126 + 2 Ma), and Zhaotun (gneiss; LA-ICPMS 128 + 5 Ma) (Wu et al. 2005), some of which, however, were previously regarded as Triassic plutons. Therefore, some geologists believed that a gap in plutonic activity ranges from c. 160 to c. 100 Ma in Korea (Kim 1996; Chough et al. 2000)

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Fig. 2. Tectonic framework of Liaodong Peninsula showing the sampling sites. Insets: A, stereo plot of 69 Yanshaninan lineations in the lower himisphere B, C, D, E, quarts c-axis preferred senses sampled from the upper part of the Archaean basement; the locations for A–E marked in the map. The contour intervals are 2%–4%–8%– 16% of 200 grains in sample B, 4%–16%– 12% of 160 grains in sample C, 2%– 4%– 6%– 8% of 170 grains in sample D, and 8%– 16% of 200 grains in sample E.

and from c. 180 to c. 140 Ma in the eastern North China Craton. However, in the Jiaodong Peninsula the emplacement age of the Linglong granite, based on SHRIMP U –Pb dating, range from 150 to 160 Ma (Miao et al. 1998; Luo et al. 1999) and that of the Guojiadian pluton from 156 to 164 Ma (Chen 1994). In the eastern Yanshan orogen

(i.e. west of the study area), the Yiwulushan granite is dated at 164 + 9.5 Ma (U –Pb SHRIMP, single zircon), with some inherited zircons yielding ages of 1991 + 16 Ma (Luo et al. 1999); these plutons are intensely deformed. In the Liaodong Peninsula four Mesozoic stages of regional deformation are identified based on

MESOZOIC TECTONICS IN EASTERN NORTH CHINA

cross-cutting relationships in the field. The first deformation is characterized by WNW –ESE- or east –west-trending thrust faults and folds in the west (Chen 1998) and north of the study area (Fig. 2) and the area near Bohai Bay. In Triassic plutons some ductile shear zones are preserved (Chen et al. 1999; Yu et al. 2002). The second deformation is characterized by NNE– SSW-trending, early and middle Jurassic, subhorizontal, ductile shear zones and folds and the 150– 160 Ma plutons. The 120 –130 Ma granites cut fabrics related to the second deformation zone (Fig. 3). At Shuiyuandi Town in the southern Liaodong Peninsula, some sheath folds (see Fig. 6), related to this thrusting event, have strong mineral lineations (see inset A in Fig. 2) and quartz c-axes (see insets B–E in Fig. 2) indicating top-to-the-NW motion. The third deformation is characterized by NNE–SSW-striking sinistral strike-slip and transtensive faults, metamorphic core complexes (Xu et al. 1987; Liu et al. 2005; Fig. 1), and rifting responsible for related Cretaceous pull-apart basins of Aptian age (c. 125 –112 Ma). The fourth Mesozoic deformation is characterized by open east –west-trending folds in the Liaodong Peninsula (Hong 1989). From the Late Cretaceous to the Early Tertiary east –west-trending thrusts and folds developed intensively in Korea (Kim 1996) and weakly in the Jiaodong and Liaodong Peninsulas. These deformation movements are based on several disconformities without precise dating constraints.

Zircon U – Pb geochronology and timing of deformation of the Dandong Granites In this section we use SHRIMP U –Pb zircon data from micaceous monzogranitic gneisses in the Dandong Granite of the Liaodong Peninsula to provide constraints on some of the tectonic questions posed above. The significance of the results is discussed in the following section.

SHRIMP results from the Dandong Granite Samples LJ023 and LJ030 are mica monzogranitic gneisses collected from Jianzhan Village of Hushan Town about 15 km north of Dandong City, and from Dongjiapu Village about 20 km SE of Xiuyan County, respectively (Fig. 2). They have a coarse granular texture and mylonitic, gneissic structure. SHRIMP age data for samples LJ023 have been published by Li et al. (2004). Here we provide similar data for sample LJ030, to determine the age of the Dandong Granite (Table 1). Zircons from sample LJ023 are euhedral, elongate, prismatic, semi-transparent to transparent. Length/width ratio of most zircons may be up to

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10:1, but is generally 5:1. Maximum and minimum zircon sizes are about 0.4  0.05 mm2 and about 0.1  0.05 mm2 in sample LJ023, respectively. Cathode luminescence (CL) and backscatter electron (BSE) images reveal that the zircons are characterized by well-developed oscillatory zoning, and zoned cores are often truncated by a zoned rim (Fig. 3). SHRIMP analytical results for zircons in sample LJ023 show that only one zircon grain (spot 30.1) on the concordia line gives a 207Pb/206Pb age of 686 + 19 Ma, the geological significance of which remains unknown. Seven concordant points of the zoned cores of different zircon grains yield a weighted mean 207Pb/206Pb age of 2105 + 21 Ma (Li et al. 2004) that we interpret as the crystallization age of major inherited zircons. All SHRIMP data in the oscillatory zoned rims and zoned cores give a younger age of 157.4 + 5.7 Ma (Li et al. 2004) that is interpreted to represent the crystallization age of this pluton and an intensive crustal reworking event in this area including the Liaonan Archaean complex reworked in the Mesozoic. This interpretation is supported by CL images of the rims that reflect magmatic zircons in spite of typical Th/U ratios. Therefore we interpret the ages as the emplacement ages of the granite. Reworked crust is mainly composed of Palaeoproterozoic plutons, revealed by a peak in the population of inherited cores of Palaeoproterozoic zircons. We recognize four morphological types of zircon within sample LJ030, which are complicated. The maximum zircon size is about 0.5  0.125 mm2 and minimum zircon size is about 0.075  0.025 mm2. SHRIMP ages for the type I zircons of this sample are scattered from 638 Ma to 127 Ma (Table 1); in particular, the age 638 + 20 Ma, 465.6 + 8.7 Ma, 206.2 + 3.8 Ma and 166.9 + 3.0 Ma are on or near the concordia line (Fig. 4). The geological significance of these ages remains unknown. However, the last two are probably related to regional crustal reworking of this area including Mesozoic reworking of the Archaean Liaonan complex (Yin & Nie 1996). It should be noted that spot 23.2 in the rim of type IV gives the same age as that of type I (Fig. 3). This leads us to suggest that the last two zircon ages represented magmatic events or remelting of the crust. The other ages were perhaps affected by later disturbance of the intense Late Yanshanian tectonism and subsequent lead loss. Two grains of type III zircons on the concordia line give a 207Pb/206Pb age of 2167 + 11 Ma to 2174 + 11 Ma, which we regard as the crystallization age of earlier Palaeoproterozoic plutons because their Th/U ratios and those of type I fall into two distinct age groups (Table 1).

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Fig. 3. Representative selection of CL zircon images from the Dandong granites (taken by G. C. Zhao in Mainz University).

Nine concordant points on the zoned cores of different type II zircon grains yield a weighted mean 207Pb/206Pb age of 1863 + 15 Ma, interpreted as another magmatic event reflected by inherited zircons of this pluton (Table 1, Fig. 4). Type IV zircons in this sample are found only in the cores of type I and they have a different zoning style from that of the peripheral zones of zircons in CL and BSE images. Their U –Pb SHRIMP ages of 2534 + 45 Ma in spot 23.1 and 2440.2 + 5.1 Ma in spot 27.1 are interpreted as the ages of different inherited zircons from Archaean and Palaeoproterozoic basement (Table 1).

Precise timing of deformation of the Dandong Granite Wang et al. (2000a, b) identified the precise timing of the first deformation event to be between 195 and 193 Ma using K –Ar and 40Ar/39Ar ages of muscovites in east –west-trending shear zones on Liaodong Peninsula. The earliest deformation event is Triassic in the Yanshan orogen (Chen 1998). Further studies integrating observations of plutonic and deformation features with geochronology have been undertaken on the granitoids around Dandong City, and especially on the Sanguliu (LA-ICPMS, 125 + 3 Ma) and Wulong plutons (LA-ICPMS, 125 + 4 Ma) (Wu et al. 2005) (Fig. 5). The new zircon ages of these plutons

range from 125 to 131 Ma (Wu et al. 2005). In this paper we have obtained Mesozoic SHRIMP ages for two plutons north and south of the Sanguliu and Wulongbei plutons, previously identified as Palaeoproterozoic. We suggest that c. 157– 158 Ma Mesozoic granites of the Dandong suite may be more abundant than previously recognized. Wu et al. (2005) also obtained a c. 165 Ma age of emplacement for so-called Pre-Sinian, deformed granites near the Wulong gold mine. In the Dandong area the c. 157–158 Ma Dandong plutons are intensely reworked into orthogneisses by subhorizontal, ductile, thrust-related shearing and subsequent NE – SE-trending folds. The Sanguliu and Wulongbei plutons cross-cut the solid-state shear foliation and folds in the Dandong orthogneisses. Thus, we deduce that the time of the second deformation of the Dandong granites is loosely defined between 157 and 130 Ma. The Linglong and Guojialing plutons in the Jiaodong Peninsula were also involved in this episode of deformation, and have yielded a muscovite 40Ar/39Ar age of 134.26 + 0.34 Ma in the Zhaoyuan–Pingdu shear zone (Lin et al. 2000). The Zhaoyuan–Pingdu shear zone cross-cuts the earlier foliation in the Linglong and Guojialing plutons, so we conclude that the precise timing of the second deformation was between 153 and 145 Ma. To the west of this area this deformation is comparable with the 161–148 Ma Chengde–Xinlong thrust and the

LJ030-1.1 LJ030-2.1 LJ030-3.1 LJ030-4.1 LJ030-4.2 LJ030-5.1 LJ030-6.1 LJ030-7.1 LJ030-8.1 LJ030-9.1 LJ030-10.1 LJ030-11.1 LJ030-12.1 LJ030-13.1 LJ030-14.1 LJ030-15.1 LJ030-16.1 LJ030-15.2 LJ030-17.1 LJ030-18.1 LJ030-19.1 LJ030-20.1

Spot

R C C C R C C C C C C C C C C C R R C C C C

1,O 1,O 2,U 3,O 1,O 2,O 1,O 1,O 2,U 2,U 2,U 2,U 2,U 3,U 2,U 3,O 1,O 3,O 2,O 1,O 2,O 2,U

1.31 0.16 – – 1.78 – 1.05 1.1 – – – 0.14 – – – – 0.97 0.11 0.45 2.38 2.45 0.01

246 627 475 360 2123 375 955 236 462 466 436 333 888 312 3101 291 561 417 892 174 158 578

41 51 250 580 161 72 224 99 229 221 242 130 526 229 751 174 115 118 146 100 77 288

Spot Textureþþ % U Th þ location cousinhood (ppm) (ppm) 206 Pb Th/ U

0.17 0.08 0.54 1.66 0.08 0.2 0.24 0.43 0.51 0.49 0.57 0.4 0.61 0.76 0.25 0.62 0.21 0.29 0.17 0.6 0.5 0.52

238

232

Table 1. SHRIMP U– Pb isotopic analyses for zircons from sample LJ030

5.94 14.2 136 124 36.7 81 26.2 3.92 138 135 123 94.4 258 107 664 103 12 132 31.6 5.15 3.33 168

Pb* (ppm)

206

Pb/ Pb age

207 206

175.2+ 4.6 144 + 280 166.9+ 3.0 189+ 69 1854 + 27 1845+ 10 2167 + 50 2265+ 47 125.0+ 2.2 2203 + 200 1442 + 24 1846+ 28 202.0+ 3.4 50+ 72 122.9+ 2.8 290 + 300 1925 + 28 1885+ 10 1879 + 28 1867+ 10 1826 + 27 1837+ 12 1834 + 28 1856+ 15 1879 + 28 1857.2 + 8.6 2170 + 32 2167+ 11 1434 + 21 1843.1 + 5.0 2218 + 32 2172+ 11 157.4+ 3.7 13 + 150 2029 + 30 2053+ 14 260.9+ 4.5 410+ 39 214.4+ 4.4 2113 + 240 158.1+ 3.9 1120 + 180 1880 + 27 1843+ 12

Pb/ U age

206 238

0.0489 + 12 0.0499 + 2.9 0.11281 + 0.55 0.1431 + 2.7 0.0424 + 7.9 0.1129 + 1.6 0.047 + 3.0 0.0521 + 13 0.11532 + 0.56 0.11421 + 0.58 0.11227 + 0.64 0.11349 + 0.81 0.11356 + 0.48 0.13521 + 0.64 0.11268 + 0.27 0.1356 + 0.63 0.0463 + 6.1 0.1267 + 0.80 0.05494 + 1.8 0.044 + 9.7 0.077 + 9.0 0.11267 + 0.64

Pb*/ Pb* + %

207 206

Pb*/ U+%

207

0.186 + 12 0.1803 + 3.5 5.184 + 1.8 7.88 + 3.8 0.1145 + 8.1 3.901 + 2.4 0.2064 + 3.5 0.138 + 13 5.535 + 1.8 5.329 + 1.8 5.07 + 1.8 5.148 + 1.9 5.298 + 1.8 7.46 + 1.8 3.872 + 1.7 7.68 + 1.8 0.158 + 6.5 6.46 + 1.9 0.3128 + 2.5 0.205 + 9.9 0.263 + 9.3 5.261 + 1.8

235

0.219 0.522 0.95 0.704 0.219 0.769 0.501 0.168 0.95 0.946 0.936 0.907 0.963 0.939 0.987 0.939 0.361 0.906 0.706 0.212 0.265 0.935

Err corr.

(Continued)

0.02755 + 2.7 0.02622 + 1.8 0.3333 + 1.7 0.4 + 2.7 0.01959 + 1.8 0.2507 + 1.9 0.03184 + 1.7 0.01925 + 2.3 0.3481 + 1.7 0.3384 + 1.7 0.3275 + 1.7 0.329 + 1.7 0.3384 + 1.7 0.4002 + 1.7 0.2492 + 1.7 0.4106 + 1.7 0.02471 + 2.3 0.3699 + 1.7 0.0413 + 1.8 0.03381 + 2.1 0.02483 + 2.5 0.3387 + 1.7

Pb*/ U+%

206 238

R R C R C R C C C R R R-C C R

2,O 2,A 4,X 2,O 1,O 1,O 2,O 4,U,X? 2,U 2,U 2,U 2,O 2,A 3,U

2.57 6.42 0.01 3.8 1.25 1.2 0.04 0.01 – 0.1 0.74 2.16 4.07 –

376 206 170 174 336 359 770 890 2094 170 246 280 154 682

270 97 93 24 211 172 65 496 594 45 103 210 192 325

U Th Textureþþ % Spot cousinhood (ppm) (ppm) locationþ 206 Pb Th/ U

0.74 0.49 0.57 0.14 0.65 0.5 0.09 0.58 0.29 0.27 0.43 0.78 1.29 0.49

238

232

29.8 13.4 70.7 3.17 9.41 6.26 153 361 624 48.6 66.2 25.1 2.61 213

Pb* (ppm)

206

Pb/ Pb age

207 206

Pb*/ Pb* + %

207 206

568.7+ 9.7 786+ 44 0.0654 + 2.1 465.6+ 8.7 434 + 130 0.0555 + 5.8 2547 + 41 2534+ 45 0.1676 + 2.7 132.1+ 4.4 382 + 230 0.0543 + 10 206.2+ 3.8 195 + 100 0.05 + 4.3 127.4+ 2.6 275 + 260 0.0447 + 11 1340 + 21 1762+ 95 0.1078 + 5.2 2495 + 34 2440.2 + 5.1 0.15855 + 0.30 1921 + 28 1812+ 22 0.1108 + 1.2 1852 + 29 1764+ 22 0.1079 + 1.2 1754 + 28 1764+ 30 0.1079 + 1.6 638 + 20 627+ 81 0.0607 + 3.7 123.0+ 3.2 86 + 490 0.0477 + 21 1994 + 29 2124+ 11 0.13191 + 0.63

Pb/ U age

206 238

Pb*/ U+%

206 238

Err corr. 0.831 + 2.8 0.0922 + 1.8 0.647 0.573 + 6.1 0.0749 + 1.9 0.315 11.2 + 3.3 0.4846 + 1.9 0.586 0.155 + 11 0.0207 + 3.4 0.317 0.224 + 4.7 0.03251 + 1.9 0.403 0.123 + 11 0.01995 + 2.0 0.184 3.43 + 5.5 0.231 + 1.7 0.311 10.33 + 1.7 0.4726 + 1.7 0.984 5.3 + 2.0 0.3471 + 1.7 0.814 4.95 + 2.2 0.3328 + 1.8 0.834 4.65 + 2.4 0.3128 + 1.8 0.745 0.87 + 5.0 0.104 + 3.3 0.662 0.127 + 21 0.01926 + 2.6 0.127 6.59 + 1.8 0.3626 + 1.7 0.937

Pb*/ U+%

207 235

C, core; R, rim; 1, Type I;, 2, Type II; 3, Type III, 4, Type IV; O, oscillatory zoned zircons; A, a large, highly luminescent core with a narrow (,m 20m), lower luminescent, oscillatory zoned rim; U, extremely low luminescent, weakly zoned zircon; X, xenocrystic core.

LJ030-21.1 LJ030-22.1 LJ030-23.1 LJ030-23.2 LJ030-24.1 LJ030-25.1 LJ030-26.1 LJ030-27.1 LJ030-28.1 LJ030-29.1 LJ030-30.1 LJ030-31.1 LJ030-32.1 LJ030-33.1

Spot

Table 1. Continued

MESOZOIC TECTONICS IN EASTERN NORTH CHINA

179

Fig. 4. Concordia plots of SHRIMP U –Pb zircon analytical results for mica monzogranitic gneiss of sample LJ030.

143 Ma Sihetang nappe (from Davis et al. 1996; Ratschbacher et al. 2000). From the Dongjiagou shear zone in the Dalian area (Fig. 2) Yin & Nie (1996) obtained muscovite 40 Ar/39Ar ages ranging from 153.1 + 1.9 Ma to 144.6 + 1.1 Ma. However, the 40Ar/39Ar ages of

deformation in the Dongjiagou shear zone and the Liaonan Archaean basement, obtained from biotite, K-feldspar and muscovite (and minor hornblende) (Yang et al. 2004) range from 109 to 114 Ma. These are identical to ages from 112.9 þ 0.4 Ma to 110.4 þ 0.4 Ma (Yin & Nie

Fig. 5. Structural map showing the Sanguliu and Wulong plutons cross-cutting the second deformation fabrics in the Dandong area.

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1996) obtained for the western detachment fault of the Liaonan metamorphic core complex. Most garnet-bearing granites, emplaced along the Dongjiagou shear zone at about 120 Ma, contain c. 1900 Ma inherited zircons (Yang et al. 2004). This relationship is similar to that of the Guojialing pluton in the Jiaodong Peninsula (Luo et al. 1999). The above relationships between dated plutons and deformation fabrics confirm that an age range from 153 to 145 Ma is most probable for the duration of the second deformation event The ages of the Yinmawanshan, Qingshan, and Zhaotun plutons are in the range 120 –130 Ma (zircon U –Pb), and are indistinguishable within error (Wu et al. 2005). According to the ages and the field investigation, the outer part of the Yinmawanshan pluton crystallized during deformation, whereas the inner part is post-deformational. The c. 120 –145 Ma marginal deformation in the plutons is different from the second penetrative deformation in the Dandong plutons. Therefore, we conclude that it was a later deformation controlled by sinistral strike-slip faulting and later extension, combined with emplacement-related structures. The sinistral strike-slip faults controlled and sometimes cross-cut the 120 –130 Ma plutons, and controlled Early Cretaceous basins with ages between 127 and 95 Ma. Therefore, the NNE– SSW strike-slip faulting (i.e. the third deformation) is constrained to an age range between 135 and 95 Ma. This result is similar to the 40Ar/39Ar age range of 128– 132 Ma for initial sinistral strike-slip faulting of the transtensive Tan-Lu fault in Anhui Province and to a biotite cooling age of 100 + 2.3 Ma for the Yilan– Yitong segment of the Tan-Lu fault in Jilin Province (Zhu et al. 2001). Several metamorphic core complexes in Eastern China formed in the overlapping areas between large-scale, sinistral strike-slip faults during the third deformation (Fig. 1). For example, the Yunmeng Shan metamorphic core complex formed at 128 –119 Ma based on U –Pb (zircon) ages on cross-cutting units and 40Ar/39Ar age spectra of hornblende and K-feldspar for cooling ages (Davis et al. 1996; Darby et al. 2006), the Yiwulushan metamorphic core complex at 126 –124 Ma (based on SHRIMP ages; Ma et al. 1999, 2000; Luo et al. 2001), the Wugongshan core complex at 145– 123 Ma (biotite 40Ar/39Ar method; Faure et al. 1996), and the Xingzi core complex at 127 + 2 Ma (U –Pb single zircon; Li et al. 2001). The Northern Orthogneiss Unit in the eastern Dabie orogen has an 40Ar/39Ar age of 140 – 120 Ma (Ratschbacher et al. 2000). The last ductile deformation age of the Dongjiagou shear zone is 114– 109 Ma according to geochronological data for hornblende, muscovite and biotite in the shear

zone (Yang et al. 2004); this can be regarded as the formation age of the Liaonan metamorphic core complex (Yin & Nie 1996).

Discussion and tectonic implication Relation between the Sulu orogen and the Jiao – Liao massif Many workers have debated whether rocks of the North China Craton, situated in the hanging wall of the Dabie – Sulu orogen, were significantly reworked by convergent and collisional processes in the Triassic. We report new SHRIMP U –Pb data that show that the c. 150– 160 Ma Linglong granite from the Jiaodong Peninsula contains many Indosinian inherited zircons of 200–250 Ma age (Miao et al. 1997, 1998; Luo et al. 1999). This event recorded in the inherited Triassic magmatic zircons is the same as that in the Dabie –Sulu orogen (Miao et al. 1998). This suggests that the Linglong granite originated from crustal felsic rocks and some inherited zircons came from granites related to the collision between the North and South China plates (Miao et al. 1997; Wang et al. 1998; Luo et al. 1999). The Linglong, Luanjiahe and Guojialing suites were derived from this early Mesozoic basement between 165 and 125 Ma, and were probably emplaced as post-collisional granitoids (Wang et al. 1998). The Linglong granite contains c. 3446 + 2 Ma to 3114 + 4 Ma inherited zircons and the Guojialing pluton many Palaeoproterozoic inherited zircons (Luo et al. 1999). This indicates that the crust in the Jiaodong Peninsula is as old as the Anshan Archaean basement in the Liaodong Peninsula, and the crust belongs to the North China Craton (Song et al. 1996) rather than the Yangtze Craton as proposed by Faure et al. (2001). To the north there are fewer inherited Triassic magmatic zircons in samples LJ030 and LJ023 from the Dandong granites than from the Linglong granite. There are three possibilities for the origin of the 640 and 690 Ma inherited zircons in samples LJ030 (Table 1) and LJ023 (data of Li et al. 2004). The first is that they come from the Yangtze Craton or the Sulu orogen. However, Hacker et al. (2000) pointed out that the effects of the Dabie –Sulu orogen did not reach that far. The 1000–800 Ma Jinning event (Li 1999) is an important marker to distinguish the North China Craton and the Yangtze Craton. Therefore, we exclude the first possibility. The second possibility is closely associated with a 680–650 Ma rifting event (Qiao 2002) recorded in 585–723 Ma diabase sills (K –Ar method; Pan et al. 2000) in the Eastern Block. Chen et al. (2002) reported

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Fig. 6. Sheath folds at Shuiyuandi Town in the Liaodong Peninsula (view towards the east; the sense of the sheath fold is toward, the west; the coin is about 2 cm in diameter).

U –Pb zircon ages of 658 –540 Ma for magmatism and metamorphism from the Huadian area. Therefore, the third possibility is that the two inherited magmatic zircons came from the Huadian area of the Jilin Province on the northern margin of the Eastern Block. The characteristics of the last two models are very different from those of the Yangtze Craton, but fit the situation of our study area. Although the precise geological implications are still unknown, our data imply that the Liaodong basement and its Mesozoic magmatism are mostly similar to the Jiaodong basement and its Mesozoic magmatism. Although the Mesozoic evolution of the Liaodong area is similar to that of the Jiaodong area, it is also closely related to the Sulu orogen, but the Jiao-Liao massif is not a part of the Yangtze Craton. Some Palaeoproterozoic inherited zircons in samples LJ023 and LJ030 indicate that the remelting was in situ. Therefore, the second possibility best matches the geological data.

Setting of remelting of continental crust and tectonic regime transition in Eastern China Although the timing of Mesozoic deformation in the eastern North China Craton is not well constrained, the north–south shortening is generally regarded to be responsible for a series of east– west-trending fold and thrust belts in the Early Mesozoic, especially the Late Triassic (Shu & An 1994; Yin & Nie 1996; Chen 1998; Chen et al. 1999; Yang et al. 2002). NNE–SSW-trending Basin-and-Range-style extensional tectonic features are commonly considered to have formed in the Late Mesozoic. However, it is uncertain how and when the transition from compression to extension occurred. Research on the transition of the tectonic regime has mainly focused on the Yinshan–Yanshan orogen

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(Zhao et al. 1994; Darby et al. 2006). Strong east – west-striking folds and thrusts that formed from the Early or mid-Triassic to Jurassic, and locally in the Cretaceous, characterize both the southern and northern margins of the North China Craton, especially the Yinshan–Yanshan orogen (Davis et al. 1998, 2001, 2002) and the Qinling orogen (Zhang et al. 1996; Meng & Zhang 2000). The time of this thrusting in the Yunmeng Shan, east of the Yinshan– Yanshan orogen, is defined at 142–143 Ma (Davis et al. 1996), possibly related to the final closure of the Mongol– Okhotsk Ocean (Yin & Nie 1996; Halim et al. 1998), or, more probably, the formation of the nearby Solonker suture (Xiao et al. 2003). In the northern part of the Sichuan Basin Early Cretaceous foreland deformation of the Qinling orogen was related to the collision between the North and South China Cratons (Zhang et al. 1996; Meng & Zhang 2000). According to palaeomagnetic data, the Siberia Craton and the North and South China Cratons were unified at c. 130 Ma (Gilder & Courbillot 1997) or from 130 to 110 Ma (Nozaka & Liu 2002) According to this south– north directed collision model of three continental blocks, we should find more east–west-trending structures of c. 130 Ma. However, in Eastern China many NE –SW- to NNE –SSW-trending thrusts and folds of Late Jurassic to Early Cretaceous age are associated with Mesozoic magmatism and deformation in the Dabie orogen, the Western Shandong Block, the Yinshan –Yanshan orogen, and the Jiao–Liao massif. Many similar events reflect similar dynamics and processes, including a widespread compression –extension transition in North and NE China. The NE– SW- or NNE–SSW-trending thrusts and folds from NW to SE in the North China Craton include the top-to-the-SSE Sihetang fold nappe in the Yunmeng Shan of Beijing (Davis et al. 1996), the top-to-the-NNW Chengde thrust in Hebei Province, the top-to-the-NW thrusts between Luanping City and Jiangchang City in western Liaoning Province, the top-to-the-NW or -west Liaonan thrusts and folds in southern Liaoning Province, the Candong thrust and the Lankao –Liaocheng thrust (buried by Tertiary strata of the Bohai Basin), the Wujing thrust and the NE –SE-trending open folds in western Shandong Province, and the top-to-the-NW or -west Xuhuai thin-skinned thrusts (Fig. 1). This north–south compressive geodynamic setting could not cause the NE –SW- to NNE –SSW-trending structural features. What geodynamic forces led to this change from east –west trending to NE–SW- to NNE–SSW-trending structures? Engebretson et al. (1985) proposed that the subduction directions of the Paleo-Pacific plate along

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the margin of eastern Eurasia changed during this interval; for example, from NW-directed convergence from 180 to 145 Ma, to NNW-directed convergence from 145 to 85 Ma, and to west-directed convergence after 85 Ma. However, the related deformation, magmatism and metamorphism in North China are less well understood, and must be analysed in the context of possibly being related to these changing subduction dynamics. We here assess the structural and geochronological evidence and consider whether the NNE–SSW- to NE– SW-trending deformation features are related to NW to NNW-directed subduction of the PaleoPacific plate in the period 180 –145 Ma or 145 – 85 Ma. In this study we identify a c. 160 Ma remelting of Palaeoproterozoic plutons in the Jiao-Liao massif. In samples LJ023 and LJ030 from near the Liaonan Neoarchaean terrane and sample LJ037 from the northern margin of the Palaeoproterozoic terrane, some magmatic, euhedral single zircons and metamorphic zircon rims around c. 2100 Ma cores are recognized, and the remelting event is constrained between 167 and 157 Ma. However, this remelting event was not previously recognized in northern or NE China. However, in SE China Du (1998) briefly mentioned such a magmatic event at 165– 137 Ma. What induced the remelting of the crust from NE to SE China? We suggest that it is not a coincidence that the remelted belt parallels the Pacific subduction zone, and that the two are related. The remelting was obviously not a consequence of intracontinental subduction (Zhang et al. 1996) in the Qinling –Dabie orogen. Four geodynamic models may resolve the Early Cretaceous remelting of the East Block of the North China Craton. Melting was of: (1) ancient lower continental crust, mostly of the North China Craton; (2) contaminated lower crust of the Yangtze crust (Zhang & Sun 2002; Zhang et al. 2002); (3) subducted slabs of the Paleo-Pacific plate (Engebreston et al. 1985); (4) a deep lithospheric wedge contaminated by crustal slabs related to the Mongol– Okhotsk Ocean, the Solonker suture (Xiao et al. 2003), and/or the Paleo-Pacific ocean (Zhou et al. 2001). However, Gao et al. (1999), Hacker et al. (2000) and Ratschbacher et al. (2000) proposed that the subducted Dabie slab was delaminated at about 190 Ma and did not affect areas farther away than 20 km north of the Dabie orogen because of the steep angle of subduction. Therefore the probability of the second model (Zhang et al. 2002) is less. The first model is also not acceptable because the geochemistry of Early Cretaceous calc-alkaline basalts and mafic dykes derived from enriched lithospheric mantle differs considerably from that of the Paleozoic and Cenozoic lithosphere of the Jiaodong Peninsula and the

Western Shandong Block (Li 2000; Zhou et al. 2001). Therefore, the setting of Early Cretaceous remelting in the Jiaodong Peninsula and the Western Shandong Block (Yang et al. 2005) is most likely to be like that of the Early Cretaceous magmatism in SE China, where it formed in a dominantly extensional environment from 140 to 90 Ma (Li 2000). That is to say, the last two models are most reasonable. Ratschbacher et al. (2000) deduced that the change from transcurrent to head-on convergent Pacific subduction in the Late Jurassic and the resultant magmatism with its onset at c. 145 Ma in Eastern China may have caused the reheating of the Hong’an Block, thus facilitating crustal extension, which is defined to be about 130–110 Ma in the Jiao-Liao massif. Although the four models mentioned above were not based on analysis of the remelting at c. 160 Ma, the Late Jurassic tectonic setting of compression related to the Palaeo-Pacific subduction was obviously different from the Early Cretaceous extension. The belt of c. 160 Ma plutons is aligned in a NE –SW direction perpendicular to the Palaeo-Pacific subduction direction, extends into the Central Asian Tectonic Zone to the north and the Yangtze Craton to the south, and shows evidence of NNE-striking thrusting. Thus, we deduce that this partial melting event can be best correlated with NW-directed subduction of the Palaeo-Pacific plate at 180–145 Ma (Engebretson et al. 1985; Du 1998). Far to the west of the Western Block of the North China Craton, the north–south trend of the Western Ordos fold–thrust belt may also suggest a link with Palaeo-Pacific subduction along the eastern margin of Asia (Darby & Ritts 2002). However, it is more likely to be a synthetic effect of eastward extrusion (Liu 1998) resulting from north– south compression between the northern and southern margins of the North China Craton (Zhang et al. 1996; Davis et al. 1996) and synchronous block faulting resulting from subduction of the Palaeo-Pacific plate. Because both the Yinshan –Yanshan orogen north of the North China Craton and the Qingling– Dabie orogen south of the North China Craton are synchronous with contraction from the Early and Middle Jurassic to Late Jurassic, two east –west-striking strike-slip faults in the north and south margins of the North China Craton, respectively, cut across the fold– thrust belts. On the other hand, the time of the collapse of the East China plateau in North China was also from 150 to 120 Ma, loosely constrained by the ages of the widespread adakites (Zhang et al. 2001, 2002; Davis, 2003). Plateau collapse probably accompanied the change in tectonic regime from compression to extension (Shao et al. 2001). Along the Taishan, Lushan and Yishan mountains

MESOZOIC TECTONICS IN EASTERN NORTH CHINA

in the western Shandong Province, the Western Shandong Block was intruded by abundant intermediate to mafic plutons with K –Ar ages of 110– 130 Ma (SBGMR 1991) resulting in formation of NW– SE-trending domino-type fault-bounded basins. A gabbro and pyroxenite with 40Ar/39Ar ages of 130 –115 Ma (Lin et al. 1996) resulted from reheating of the Western Shandong Block. In the Dabie orogen the Northern Orthogneiss Unit suite has 40 Ar/39Ar ages of 130 –120 Ma, obtained from hornblende and biotite (Ratschbacher et al. 2000). However, Li (2000) proposed that the 140 –100 Ma peak stage of coeval, intensive volcanism and magmatism may have been generated in response to Izanagi Plate subduction in SE China. Maruyama et al. (1997) suggested that the Palaeo-Pacific plate subducted at a fast velocity (greater than 200 mm a21) at a low angle in the period 145 –85 Ma. Because the Eastern Eurasian continental margin extended from south to north at that time, NW-directed subduction of the Palaeo-Pacific plate means oblique subduction at 145 –85 Ma. This oblique subduction led to intense shearing of the East Asian continental margin, with strain partitioned into zones of contractional strain and zones of strike-slip shearing (Maruyama et al. 1989). In this case, large-scale left-lateral faults formed and developed along the weak zone of the island arc and aided upward emplacement of magma and formation of metamorphic core complexes. However, to the south of the study area, this change from compression to extension in SE China was at about 120 –100 Ma, later than in NE China. This event resulted in thinning of the crust and lithosphere, upwelling of the asthenosphere, and basaltic magma underplating in SE China (Liu et al. 1990; Xu et al. 1999; Shu et al. 2000b). In a similar scanario, Bradley et al. (2003) and Kusky et al. (2003) related diachronous changes along the Pacific rim in Alaska to passage of a triple junction, and subduction of different plates before and after triple junction migration with different convergence vectors before and after ridge subduction. In summary, we propose that a remelting event at c. 160 Ma and a compressive deformation event from 153 to 145 Ma marked a change from Triassic compression to Early Cretaceous extension in the Liaodong Peninsula, and across North China. This remelting event is well correlated with NW-directed subduction of the Palaeo-Pacific plate.

Implications for subduction of the Pacific plate beneath the Eurasian plate The east–west-trending deformation at 200– 190 Ma in North China, especially in the Liaodong Peninsula (Chen et al. 1999; Wang et al. 2000a),

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matches the collision of the North and South China plates and 210–180 Ma subduction of the Farallon plate interpreted by Maruyama et al. (1997) and based on geology and paleomagnetic data. Mesozoic magmatism, especially from 145 to 90 Ma, related to the Pacific subduction has been studied more in SE China (e.g. Li 2000), than in NE China (e.g. Lin et al. 1998) and the northern margin of the North China Craton (e.g. Zhou et al. 2001). However, there is little understanding about the relation between subduction of the Pacific and Izanagi plates and Mesozoic magmatism between the above-mentioned three areas. Ratschbacher et al. (2000), summarizing data from Chinese metamorphic core complexes from Mongolia to SE China, concluded that the tectonic setting of Early Cretaceous (especially 147– 110 Ma) deformation was extensional in Eastern and NE China. Wang et al. (2002) proposed that two stages of volcanism occurred during this time. This was considered to result variously from Pacific back-arc extension, north–south compression between the Mongol–Okhotsk and Bangong –Sanjiang sutures, and related eastward tectonic escape of the North China and South China Blocks. However, all the metamorphic core complexes mentioned above are aligned in a NNE –SSW direction parallel to the Pacific subduction zone. Thus, the mid-Cretaceous tectonic setting was dominated by Pacific subduction (Ratschbacher et al. 2000). Cretaceous magmatism in SE China also occurred in a dominantly extensional environment, especially in continental back-arc and postcollisional extensional settings from 140 to 90 Ma according to A-type granitic and within-plate basaltic geochemistry (Shu et al. 1998; Li 2000). Therefore, we suggest that the transition from compression to extension occurred over a time range from 180 to 145 Ma. The transition time was previously suggested to be between 120 and 90 Ma by Maruyama et al. (1997). The remelting in the Jiao-Liao massif suggests that the monzogranites underwent intensive Mesozoic reworking. At 180–120 Ma there was an Andean-type continental margin in Eastern China (Li 2000; Li et al. 2001). Furthermore, our data and those of Li (2000) and Li et al. (2001) indicate that the time of NNE –SSW striking contractional folds and faults that was closely related to the Pacific subduction was from 160 to 145 Ma, supporting an extensional time mainly from 145 to 100 Ma. During this time Eastern China was affected by coeval wrench faulting, voluminous volcanism and extensional basin formation. In the period 100–90 Ma another intense compressive deformation occurred along the continental margin of the western Pacific, especially in SW Japan, South Korea, SE China, Taiwan and the western

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Philippines. This was possibly related to accretion and collision between terranes on the oceanic plate and the continental margin of the Eurasian plate (Charvet et al. 1985, 1999; Ichikawa et al. 1990; Zhao et al. 2007). This process gave rise to the formation of ophiolite me´langes, high-pressure blueschists, paired metamorphic belts, collisiontype granites and ductile shear zones along the suture zone (Charvet et al. 1990; Shu et al. 2000b). This study was funded by China NSFC grants 40472098 and 40002015, grant KZCX1-07 of the Chinese Academy of Science, and Hong Kong RGC grant HKU7090/01P, and was also supported by US NSF grant EAR-02-07886 awarded to T. M. K. We thank our colleagues Hao Defeng, Luo Yan and Xia Xiaoping for their fieldwork, and M. Brown, S. Maruyama, Wang Tao and Moonsup Cho for sending some important papers. SHRIMP dating was supervised by Wan Yusheng, Song Biao and Liu Dunyi and Jian Ping at the Beijing SHRIMP Centre their work is appreciated. We thank two anonymous reviewers and B. Windley for comments that helped to improve the manuscript.

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Deep crustal structures of the Yanshan intracontinental orogeny: a comparison with pericontinental and intercontinental orogenies J. SHAO1, G. HE1 & L. ZHANG2 1

Faculty of Earth and Space Sciences, Peking University, Beijing 100871, China 2

Inner Mongolia Institute of Mineral Experimentation, Huhhot 010020, China (e-mail: [email protected])

Abstract: The Mesozoic Yanshan intracontinental orogenic belt is developed in a weak zone on the northern margin of the North China Craton. We classify the Yanshan as an intracontinental orogen, as opposed to an intercontinental or pericontinental orogen, because of several geodynamic factors that differ from those of typical intercontinental orogens. These are: (1) reactivated tectonic activity on the lithospheric-scale faults since 1.8 Ga; (2) differential uplift of blocks during the Mesozoic, especially the great change of the elevation–subsidence framework at about 140 Ma, which induced building of the Yanshan Mountain and formation of the Cretaceous basins; (3) polycyclic evolutionary processes of active rift basins recorded by the Mesozoic volcanic and sedimentary strata; (4) Mesozoic tectonic regime reversion in the eastern North China Craton, demonstrated by petrological and geochemical evidence of crust– mantle interaction and supported by research on lithosphere thinning and ancient heat flow. Among these factors, the acute tectonic events of the lithosphere and evolution of active rift basins are directly affected by the upper mantle. The differential elevation of faulted blocks directly reflects the deep tectonic processes. Inherited features of the tectonic activities are characteristic of intracontinental orogenies and occur throughout all stages of the Yanshan movement but with different features. The deep controlling factors changed with time and combined with each other. The depth of tectonomagmatic activity became progressively shallower from 180 Ma to 130 Ma and then became deep again after 120 Ma. In the same period, extensional deformation was predominant in this area. The most important geodynamic characteristic of the Yanshan intracontinental orogeny is the tectonic regime inversion. Intracontinental orogeny is the response of the upper crust to the dramatic movements of the deep lithosphere. Its essential tectonic processes are thickening and subsequent thinning of the crust and lithosphere, which induces not only deformation of the rocks but also uplift of the mountain chain. Magmatic underplating is the most important factor of the Yanshan orogeny.

The ‘Yanshan movement’, named from the Yanshan area, has attracted the attention of Chinese and international geologists since the 1920s (Wong 1927) and was considered as a special or non-typical orogen because it developed in a stable craton; the North China Craton. Workers have variously named it the ‘Yanshan platform –fold belt’ (Ma et al. 1961), the ‘composite orogen caused by ocean subduction and craton resistance’ (Deng et al. 1996), the ‘North China intracontinental orogen’ (Ge 1989), the ‘Yanshan style intraplate orogen’ (Song 1999) and the ‘cratonic orogen’ (Che et al. 2002), since the theory of plate tectonics became established. There are various hypotheses about the dynamics of the Yanshan orogen. Some workers suggested that it was the result of ‘faultblock mountain-building’ (Zhang 1984) and ‘platform reactivation’ (Chen 1960), emphasizing that the vertical crustal movements control large-scale differential elevation and related magmatism.

During the 1990s, some researchers, in the light of plate tectonics, attempted to seek power sources of the Yanshan deformation from the plates adjacent to North China (Davis et al. 1996, 2001; Yin & Nie 1996). Yin & Nie suggested that the Yanshanian intra-arc fold and thrust belt in North China might have been generated in the Mesozoic by (1) the southward subduction of the Mongol–Okhotsk plate beneath North China or (2) the collision between North China and South China, or (3) the subduction of the Pacific plate beneath eastern Asia, or (4) a combination of the three. No definite conclusion was drawn. Recently, after recognition of the fact that Mesozoic lithospheric thinning occurred and deep geological activity was coupled with shallow tectonic movements in the North China Craton, the concept of the orogeny was also extended with studies of intracontinental orogenies. The present authors, impressed by the dominant role of vertical differential movements and volcanism in

From: ZHAI , M.-G., WINDLEY , B. F., KUSKY , T. M. & MENG , Q. R. (eds) Mesozoic Sub-Continental Lithospheric Thinning Under Eastern Asia. Geological Society, London, Special Publications, 280, 189– 200. DOI: 10.1144/SP280.9 0305-8719/07/$15 # The Geological Society of London 2007.

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the formation of the mountains and basins in this region, conclude that the Yanshan Mountains should be a useful example for exploring how the deep dynamic factors control an intracontinental orogeny. Orogens can be divided into three types, (intracontinental, pericontinental and intercontinental) according to their tectonic positions. The last two types form during rifting and reassembly of the continents through plate tectonics, whereas the first type is mainly controlled by deep geodynamics. This paper focuses on the geology of the Yanshan intracontinental orogen in the North China Craton to discuss the deep geodynamics of intracontinental orogens.

Inherited tectonic activities The Mesozoic Yanshan orogeny has obvious inheritance features, which are critical to understanding the evolution and nature of the Yanshan intracontinental orogen. The Mesozoic Yanshan orogenic belt is located south of the Zhangjiakou–Luanping –Chengde – Beipiao belt in the northern part of the North China Craton, including the north Taihangshan Mountains, Jundushan Mountains, and Yanshan Mountains. It

covers the same area as the ‘Yanshan platformal fold belt’ (Ren et al. 1980), and as the ‘Central Zone’ between the East and West blocks of North China as proposed by Zhao et al. (1999) (Fig. 1). After assembly of the North China Craton between 2.5 and 1.8 Ga (Zhao et al. 1999; Kusky & Li 2003; Kusky et al. 2007), the craton started to break up again after 1800 Ma (Zhai et al. 2000), with development of anorogenic magmatism in the Central Zone, including dyke swarms at 1850–1700 Ma (Li, J. et al. 2001; Shao et al. 2002). The magnetic fabrics of the dyke swarms show that the magmas flowed toward the NNW or NW (Hou & Qian 1992), reflecting the tectonic stress field during the Early Proterozoic and indicating that the Central Zone was an important passage for mantle-derived magmas. The Yanliao aulacogen (also called the Yanliao subsidence zone) strikes ENE–WSW, with the maximum thickness of Middle Proterozoic (1800–1000 Ma) sedimentary strata being more than 10 000 m. The Changzhougou and Dahongyu Formations consist of sedimentary rocks and a large amount of intermediate–basic volcanic rocks, reflecting an extensional setting at 1.8 Ga (Kusky & Li 2003). The Beijing–Chengde–Beipiao fault, marking the NW boundary of the aulacogen, is a southeastward steeply dipping synsedimentary fault

Fig. 1. Distribution of basement faults and Mesozoic strata in the North China Craton.

DEEP STRUCTURES OF YANSHAN OROGEN

(Wu & Meng 1998), which is also called the basement fault belt (Fig. 1). It is consistent with the eastern boundary of the Central Zone at the end of the Early Proterozoic, showing its inheritance. This paper agrees with the point that the intracontinental orogeny in the North China Craton began from the Indo-Sinian Period (about 230 Ma), which was the prelude to the Yanshan movement. The main orogenic stage was from the Early Jurassic to the Late Cretaceous (Cui et al. 2000). The Early Jurassic strata are distributed along the basement fault belt as mentioned above, whereas the Middle– Late Jurassic strata are distributed more extensively. However, they are all controlled by the basement fault and were the result of crustal extension (Gao 1957; Ma & Liu 1986) (Fig. 1). All these facts show that the ancient (with ages to 1800 Ma) lithosphere faults in the basement of the North China Craton, forming previous tectonically weak zones, were reactivated during the early stage of the Yanshan movement and controlled the distribution of the Jurassic faulted basins. The faults obviously controlled the distribution of the Yanshan Belt by the end of the Late Jurassic, indicating inheritance of the intracontinental orogeny.

Differential uplift of blocks The North China Craton was stable from the Neoproterozoic (1000 Ma) to the end of the Palaeozoic, with relatively weak tectonomagmatic activity. From the Late Triassic, the tectonic setting of the North China Craton changed from the Tethys tectonic domain to a new tectonic framework, with uplift in the east and sagging in the west. At the same time, the NNE –SSW-trending structures formed (Shao et al. 2000a). The adjustment of the deep structures induced differential surface elevation, which locally resulted in strong denudation, such as the denudation of Middle to Late Triassic strata of up to 3122 m and 5354 m in the east of the North China Craton (Fu et al. 2005). It is notable that there is no angular unconformity between the Triassic and Permian formations in most areas of the North China Craton (Bureau of Geology and Mineral Resources of Hebei Province 1989). Local faulted basins and box synclines were formed in some areas in settings of differential elevation in the Early Mesozoic. The Yanshan orogen was in a relatively stable extensional environment during the Early Jurassic with little difference in landforms on either side of the orogen, as shown by the similarity of coal-bearing strata of the Xiahuayuan Formation. During the Middle–Late Jurassic, extensive reactivation of the basement faults induced stronger differential topographic activities and dramatic changes in the

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facies and thickness of strata. For example, in the Tiaojishan box fault basin (Shao et al. 2003a), the compositions of conglomerates from the volcanic clastic rocks in the Tiaojishan Formation vary on the two limbs of the syncline because of erosion of different basement rocks. The thickness of the conglomerate-bearing rocks is about 300 m, covered by 837 m of volcano-clastic rocks and lavas with local accumulation of volcanic mud flows. The gradient of the thickness variation is about 400 m km21 from the margin to the centre of the basin, indicating that the accumulation is controlled by the differential movement of the faulted blocks. The dip angle of the strata in the fault basin decreases gradually with weakening of the synsedimentary fault activities in the later period. Although distribution of the Late Jurassic volcano-sedimentary strata is wider than that of the Middle–Early Jurassic strata, this distribution is still controlled by the basement faults of the craton (Fig. 1). The boundary of the Jurassic and Cretaceous (140 Ma) was a period of no sedimentation in the Yanshan Belt. Elevation of the faulted blocks significantly changed in the Yanshan area, with the Triassic– Jurassic basins being uplifted and the previous uplifted areas on both sides partly subsiding. Inclination angles of the Jurassic and the underlying Meso-Neoproterozoic beds in the uplifted areas are relatively gentle, suggesting uplift of the faulted blocks. On both sides of the uplifted areas there are Early Cretaceous fault basins mainly filled with volcanic rocks. On the NW side are the Zhangjiakou– Luanping Cretaceous basins. In these basins, lower volcanic lavas and upper lake–marsh facies sedimentary rocks, which consist of oil shale and Rehe biota, together make up the Early Cretaceous sequence, which is several kilometres thick and overlaps the Archaean metamorphic rocks (Bureau of Geology and Mineral Resources of Hebei Province 1989) (Fig. 2). On the SE side of the North China basin, the Late Mesozoic rocks overlap over the Meso-Neoproterozoic rocks. For example, the Beijing graben on the north side of the North China basin (Zhang, H. 1999) was controlled by the Babaoshan fault. The fault experienced two successive tectonic events. The early one was a thrusting event at the end of the Late Jurassic, shown by the Mesoproterozoic and Late Jurassic strata thrust onto Carboniferous and Permian strata. The later one was the accumulation of Cretaceous and Tertiary continental clastic and volcanic sediments with thicknesses of more than 1 km in the normal-fault-controlled basins. In the Daxing uplifted area, SE of the Beijing graben, Quaternary strata directly overlie Cambrian rocks (Fig. 2). There was a compressional event at the Jurassic– Cretaceous boundary, named the Late Jurassic ‘thrust–nappe event’ (Zheng et al. 2000). The

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Fig. 2. Sections showing the Luanping fault basin and the Beijing graben in Early Cretaceous time.

youngest strata involved in this event are the Tiaojishan and Houcheng Formations; the Early Cretaceous Zhangjiakou Formation was not involved (Shao et al. 2003b). Nappe structures of this period commonly occur in many areas, such as Yixian, Xuanhua, Huailai, Changping, Tanghekou, Chengde and Luanping (Fig. 3). The compressional event shows three features. (1) Thrust or nappe faults are dominant without large-scale strongly compressive folds (most of the folds were accompanied or induced by faults). Therefore, the angular unconformity between the Jurassic and Cretaceous strata at Luanping, Chengde, and other localities is not significant (less than 308 in general). (2) The thrust or nappe faults are not continuous and are of limited extent. The displacement distances on the Nandazhai, Yunxialing, and Huangshandian faults in SW Beijing are 24 km, 5.11 km and 2.75–1.25 km, respectively, based on the equilibrium profile

calculation (Shan et al. 1991). The displacement distance of the Chengde thrust fault is 40–45 km, based on calculation of the flattened thrust–fault plane (Zheng et al. 2000). Most of those faults, located in the western and northern parts of the upwarp belt of the Meso-Neoproterozoic basement and thrust or nappe, are perpendicular to the arc-like upwarp belt; for example, the Xilin–Yaoshunkou fault of Yixian county thrusts toward the 2908, the Nandazhai fault thrusts toward 3108, the faults including the Jimingshan fault at Xuanhua, the Changping fault, and the Xinglong fault thrust toward the 3308, and the Chengde fault thrust to 3408. (3) The basement Proterozoic strata are often incorporated into the thrust–nappes, forming so-called ‘thick-skinned structures’ and resulting in steeper dip angles at the back of the faults (Zhang & Song 1997). According to the above features, we suggest that tectonism of this period should be associated with changes of the

Fig. 3. Sections showing Late Jurassic reverse thrust–nappe structures of the Yanshan Mountains.

DEEP STRUCTURES OF YANSHAN OROGEN

elevation in these areas, and the northwestward thrust is a reactivation of the previous steeply dipping faults, which transform into low-angle faults and accompanying folds near the surface. Besides the vertical differential movement, we do not rule out a possible NW–SE-trending horizontal compression, which is possibly related to the large-scale tectonic setting and needs to be studied further. From 130 to 105 Ma, uplift and magmatism became strong in the Yanshan area, and a series of metamorphic core complex or diapiric thermal domes formed in the places such as Zhang-Xuan, Yunmengshan, Harqin and Fangshan (Fig. 1; Davis et al. 1996). After the Late Cretaceous, volcanic activity stopped and tectonic activity became weak in the Yanshan area; this is also confirmed by fissiontrack analyses showing that this area entered a slow cooling stage from 80 Ma (Wu & Wu 2003; Shao et al. 2005). All of these features mark the end of the Yanshan orogeny. In summary, the entire orogenic process of the Mesozoic Yanshan Belt, from the initial rifting to upwelling and mountain-building, was accompanied by a series of magmatic activities and tectonic events with differential elevations across the zone. The term ‘mountain building process’ (Anonymous 1986) includes not only the deformation process but also upwelling of the mountain. Therefore, it is reasonable that workers proposed that end of the Yanshan orogenic process was in the Late Cretaceous.

Active rift basins in the Cretaceous Fault and volcanic activity are widespread in the Yanshan intracontinental orogen, expressed by polycyclic evolution of active rift basins. Active rift basins were caused by upwelling of mantle materials and characterized by volcanism prior to sedimentation (Burke & Dewey 1973). Most workers have divided the Jurassic – Cretaceous strata into four volcano-sedimentary cycles: the Nandaling, Tiaojishan, Zhangjiakou, and Dabeigou cycles (Table 1). It is evident that each cycle, except the Zhangjiakou cycle, began with volcanic activity and ended with lake –marsh facies sedimentation. The basins experienced a complete cycle ranging from fault-controlled basins to downwarping, and then a hiatus in sedimentation during which compression occurred, involving deformation and uplift, and forming unconformities, disconformities and angular unconformities.

Nandaling cycle The Nandaling cycle, the prologue to the Yanshan orogeny, consists of alkali basalts, indicating rifting. The thickness of alkali basalt and minor

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andesite ranges from 15 to 767 m at Panjianzigou west of Beijing, where there are four to five sedimentary horizons intercalated with the alkali basalt, reflecting a syndepositional rift environment that controlled the volcanism in the Early Jurassic. The thickness of basalts decreases away from the basement faults. The basalts of the Nandaling Formation were derived from an enriched lithosphere mantle, contaminated by lower crustal material of the North China Craton (Li et al. 2004). The Nandaling basalts are continental intraplate basalts according to the tectonic setting identification diagram for elements Th, Nb, Ta and Hf, and are extensional or rift basalts (Shao & Zhang 2004). The K –Ar isotopic age of the basalts is 177–198 Ma (Sun et al. 1994). The volcanic basins (of the Xiahuayuan Formation) were transformed into coal-bearing lake – marsh basins at the end of the Early Jurassic, with 1000 m of sedimentary strata and 100–500 m of coal seams. The Middle Jurassic Jiulong Formation consists of a suite of purplish red, river–lake facies sediments with variable sedimentary rhythms. The upper layers, bearing marl lenses, mud nodules and pyrite, are fine-grained and are typical lake facies with horizontal lamination. Sandstones with cross-bedding increase in the upper part, indicating that the water became shallower and the sediment changed to riparian facies deposits. The volcanic materials in the sediments indicate the start of large-scale volcanism. From the Xiahuayuan Formation to the Jiulongshan Formation, the sedimentary environment changed gradually from reducing to oxidizing, from humid to dry, and from stable to active tectonics. Thus, we include the Jiulongshan Formation in the Nandaling cycle.

Tiaojishan cycle Volcanic rocks in the Tiaojishan Formation of the Tiaojishan cycle are trachytic rocks and high-K rhyolites. There are four main views about their origin and tectonic setting; that is, partial melting of an upper mantle source (Li et al. 1995), mixing of magma from crust and mantle (Li 1993; Sun et al. 1995) partial melting of lower crust (Bao et al. 1995), and partial melting of basaltic rocks that formed by basaltic magma underplating (Li, W. et al. 2001). These different views have the common feature that the Late Jurassic volcanic magma was derived from a deep level with strong interaction between the crust and mantle. There are extensive spatial variations in the lithological and lithofacies features of the Houcheng Formation, including the presence or absence of weak or strong volcanism. Recently some workers have proposed that there are bimodal volcanic rocks, based on the assemblage of basalts and rhyolite from south

Qingshilazi Huajiying

Nandian

Formation

Intermediate basic volcanic – sedimentary rocks, basalts at bottom

River– lake facies, oil shale or coal

Volcano-sedimentary formation

Magmatic origin

119 – 123 (K– Ar)*

Age (Ma)

6.22

Eo

III1 Zhangjiakou cycle, stair of Early Cretaceous

Strong high-K rhyolitic and trachytic magmatism

Pyroclastic volcanical rocks and intermediate lava

Zhangjiakou

Baiqi

High-K acidic magmas from partially melted lower crust with initial 87Sr/86Sr ratios of 0.7064– 0.7073

41 – 100

1362 (zircon U – Pb) 138 – 139 (WR Rb – Sr) 3

45 – 86

130 – 1321 (zircon U – Pb)

Xiguayuan Dabeigou Uplift, formation of metamorphic core complex, and intrusion of basic dyke swarms and granites (130 – 120 Ma)

III2 Dabeigou cycle, end of Early Cretaceous

Episode

Table 1. Mesozoic sedimentary formation, and volcanic and tectonic activity in the Yanshan area

12 – 14

35 – 50

,10

69.0

t

70.2

89.8

66.5

4.09

DI

Indices of volcanism

3.77

4.90

3.34

v(K2O)

Large-scale magmatism after volcanic eruption

Activity diminishing, fault basin changing into downwarped basin

Tectonism

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River– lake facies formation with interbedded bimodal volcanic rocks K-rich trachyandesite basaltic volcanism

Houcheng

Tiaojishan

Mixture of two magmas from upper mantle and lower crust with (87Sr/86Sr); of 0.7065– 0.7073 148 (Ar– Ar)

Nandaling

Middle – Early Jurassic Late Triassic

Continental alkali basalts River-lake facies

River-lake facies River-lake facies bearing coal

Cumulate, granulite, alkaline rocks

Primary magma produced by partial melting of upper mantle

228 – 220 (Ar – Ar)9

230(zircon U – Pb)8

177 – 198 (K – Ar)7 6.9

10 – 17

16 3.6

42.4

63.4

2.54

2.74

3.4

Large-scale surface uplift and sedimentation of fault basins, underplating at deep level

Fault basin changing into coal-bearing downwarped basin

Volcanic fault basin changing into downwarped basin

1 Zhang et al. (2005); 2Niu et al. (2003); 3Shao et al. (2003b); 4Shao et al. (2003a); 5Zhao et al. (2004); 6Davis et al. (2001); 7Sun et al. (1994); 8Hu et al. (2005); 9Shao et al. (1999). *Obtained in the present study. Eo, explosivity index; t, Gottini index (Li 1993); DI, differentiation index; WR, whole rock.

Xingshikou

Jiulongshan Xiahuayuan

I Nandaling cycle

0þ

38.8– 41.96

157 (U– Pb)5 6

50 – 70

145 (K– Ar)4

Interruption of sediment, local angular discordance, differential elevation (c. 160 Ma)

II Tiaojishan cycle, Late Jurassic

Differential elevation, interruption and denudation of sediment, angular discordance, thrust– nappe structure (to 140 Ma)

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Hunyuan, Beishuiquan in Yuxian, Chuncun and Pangjiafang in Xuanhua, Shuichang in Chicheng, and at Baihepu in Yanqing, and therefore suggested an extensional setting for the volcanism (Shao et al. 2003a; Wang & Ji 2004). Further study is needed to determine the precise tectonic setting of the volcanism. The conglomerates at the base of the Houcheng Formation show features of strong fault subsidence in the early stage of basin formation. The depositional setting of alluvial fans and braided rivers occurs in such a tectonic environment. The finegrained sediments in the middle and upper parts of the Houcheng Formation overlap onto the margins of the basins, indicating a continuous growth of the basins. The sedimentary environment changes from alluvial fans and braided rivers to meander-lakes, and the fault basins are converted into downwarped basins, indicating that the degree of extension decreased with weak volcanic activity. Thus, evolution of volcanism and sedimentation of the Tiaojishan and Houcheng Formations reflects a complete active rift process.

Zhangjiakou and Dabeigou cycles The thickness of Mesozoic volcanic rocks in the fault basins is up to 15 000 m, of which the volcanic rocks of the Zhangjiakou cycle occupy about 47%. The Zhangjiakou cycle is mainly composed of acidic volcanic rocks, including rhyolite and trachyte (Shao et al. 2003b). It is not really an independent cycle, and it is better to combine the Zhangjiakou with the Dabeigou cycle (Table 1). The intermediate –basic volcanism became significantly weaker in the early part of the Dabeigou cycle. A suite of trachybasalts and trachyandesites in the Donglanggou Formation occurs west of Beijing, which corresponds to the Dabeigou Formation. In the later part of the Dabeigou cycle, deposition river –lake facies formations occurred, showing that downwarping was dominant at the end of the Mesozoic. From the features of volcanism of each cycle (Table 1), explosivity index (Eo ¼ volcanic clastic rocks/(volcanic clastic rocks þ lavas)  100%) is directly related to penetrability of the crust. It can be seen that the volcanism changed with time from effusive facies to explosive facies as a result of the closure of various fissures and the build-up of volatile components. Finally, the volcanic rocks were replaced by plutonic rocks, with the transitional peak time at around 130 Ma. The Gottini index t ¼ [v(Al2O3) 2 v(Na2O)]/v(TiO2), in which v denotes mass percentage, extensively used to identify the tectonic background of a volcanic rock (Rittmann 1970), is directly proportional to v(Al2O3) and inversely proportional to both

v(Na2O) and v(TiO2), reflecting the degree of crustal contamination with variable K2O contents. Increase of the differentiation index (DI ¼ q þ or þ ab þ ne þ lc þ kp) indicates increasing magma differentiation, which would result in the magma being transformed to a residual magma enriched in alkalis and aluminium silicate. From the abovementioned four indices, it is evident that the degree of crustal permeability and the activity of the crust decrease gradually from the Nandaling Formation to the Zhangjiakou Formation. The occurrence of intermediate –basic volcanic rocks in the Dabeigou cycle at 120 Ma indicates that the magma source became deeper again at this time.

Evidence for strong tectonic disturbance of the lithosphere The Mesozoic Yanshan intracontinental orogen occurs on a stable craton, hence the differential elevation of the faulted blocks and the tectonic evolution of the basins must be related to deep geodynamic processes in the crust or lithosphere. The North China Craton was relatively stable from the Meso-Neoproterozoic to the Early Palaeozoic, with eruption of diamond-bearing kimberlites. The lithosphere had a thickness of about 200 km (Lu & Zheng 1996). Recently, results of calculations of ancient heat flux show that the thickness of the thermal lithosphere was 135–148 km in the Early Mesozoic (T2 – 3) and 50 –55 km in the Late Mesozoic (J3 – K1), respectively. The denudation is 3 km and 2 km, the heat flux of the crust is 24.5 mW m22 and 23.1 mW m22, and the heat flux of the mantle is 25.5 mW m22 and 56.9 mW m22, respectively (Fu 2003; Fu et al. 2005). This indicates that (1) T2 – 3 and J3 –K1 are the two periods of the Mesozoic when crustal differential elevation and geothermal activity become strongest in the North China Craton, and (2) lithospheric thinning in North China began in the Early Mesozoic. Late Triassic rocks, including accumulated rocks, granulites, alkaline rocks, A-type granites and dyke swarms, indicate strong magmatic underplating (Mu & Yan 1992; Shao et al. 1999, 2000b; Yan et al. 1999; Han et al. 2004). The peak period of the dyke-swarm intrusion is about 120 Ma, including the bimodal dyke swarms in northern Beijing. These dykes extend for several kilometres from Nankou northward through Juyongguan into the Dahaituo granite, and southeastward into the Quaternary plain, striking perpendicular to and intersecting with the abovementioned basement fault belt (Fig. 1). The dyke swarm has (87Sr/86Sr)i ¼ 0.7055 – 0.7059 and 1Nd(t) ¼ 235 to 210, indicating that the magmas

DEEP STRUCTURES OF YANSHAN OROGEN

originated from an enriched mantle (EM; Shao et al. 2001). The contemporaneous Duijiuyu granite at Changping near the dyke swarm has a Rb – Sr age of 128 Ma and an initial 87Sr/86Sr value of 0.7041 (Sun et al. 1992). Thus, it indicates that the basement fault was the passageway for the mantlederived magma and the NW–SE-striking extensional fault was the site of magma crystallization. Recently, many data have been provided concerning Mesozoic crust –mantle interaction and lithosphere thinning in the North China Craton, which provide evidence for a strong lithosphere disturbance during the Mesozoic, although different views exist on its dynamic mechanism (Deng et al. 1996; Fan & Liu 1996; Gao et al. 1998; Lu et al. 2000; Wu et al. 2000).

Summary Based on the present evidence, the Yanshan orogen was controlled by deep geodynamic processes related to the upper mantle. Lithospheric transformation, active rift basins, and differential elevation of faulted blocks reflect deformation at different crustal levels. The former two features are the direct result of deep geodynamics, and the last feature is a response to deep geodynamics. Inherited (reactivated) tectonic activity occurred throughout all stages of the Yanshan movement but with different manifestations. The deep geodynamic factors varied with time and are superimposed on each other, whereas the depth of tectonomagmatism became progressively shallower. The Mesozoic Yanshan orogen developed in a weak zone of the North China Craton (i.e. an ancient aulacogen). Tectonic deformation was mainly expressed as differential elevation of the faulted blocks involved in its basement, which resulted in the formation of open or box folds. There was no regional metamorphism. The Yanshan area has experienced multiple-stage evolution of active rift basins accompanied by later extensional deformation, with diapiric thermal domes, metamorphic core complexes and the emplacement of dyke swarms. The crust became thin before being thickened prior to final lithospheric thinning. Magmatism involved interaction between crust and mantle, as shown by mixing of magmas derived from mantle and crust or anatexis of the lower crust. The Yanshan orogen experienced a short stage of compressional deformation at the Jurassic –Cretaceous boundary (about 140 Ma), which defined the framework of the orogen. The orogeny ended with a relatively quiet extensional period in the Middle–Late Cretaceous. The differences between intracontinental orogenies and pericontinental and intercontinental

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orogenies can be summarized from the Mesozoic evolutionary history of the Yanshan orogen, as follows. 1. Pericontinental or intercontinental orogenies are one of the stages of the Wilson cycle of the tectonic evolution of plates. Orogeny is induced either by subduction or by continental collision and is always accompanied by large-scale horizontal convergence and consequent lithosphere thickening. However, the Yanshan orogen developed in the stable North China Craton and the Mesozoic orogenic process began with lithosphere thinning. The remote influences of the adjacent plates around the North China Craton may have affected the Yanshan orogeny, but lithospheric-level geodynamics of the North China Craton directly controlled the evolution of the Yanshan orogen. 2. For an intracontinental orogen, the deep tectonothermal state is the most important factor that affects the structural adjustment during lithospheric thinning. It lowers the strength of rocks locally, which induces the reactivation of basement faults to play a major role. As a result, the Mesozoic orogenic process is mainly shown by volcanic activity from fissure eruptions to planar flood events and polycyclic evolution of active rift basins. Therefore, intracontinental orogenic belts on cratons are characterized by inheritance and polycyclic events, whereas pericontinental and intercontinental orogens are developed on newly accreted continental crusts without inheritance and polycyclic events. The foreland basins are different from the inter-mountain active rift basins in the Yanshan intracontinental orogenic belt. 3. Because the driving force comes from depth, vertical heterogeneity results in differential vertical movement of faulted blocks and formation of fault basins, which are the predominant forms in intracontinental orogens. Volcano-sedimentary basins are developed symmetrically on both sides of the uplifting basement. The deformation mechanism is relatively complex, and extension and contraction exist at the same time. The deformation mainly consists of thick-skinned structures in which the basement is involved. However, a pericontinental subduction orogen or an intercontinental collision orogen is related to continentalcrustal thickening caused by horizontal convergence, with development of thin-skinned structures such as strong folds or nappes. 4. Magmatism related to crust –mantle interaction is one of the important characteristics of intracontinental orogenic belts. The continental crust mainly grows vertically, and mantle materials play an important role in crustal growth. The Sr and Nd isotopic compositions of the igneous rocks always indicate an enriched Mesozoic lithospheric mantle beneath the North China Craton.

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5. The compositions of volcanic rocks reflect the transformation of the crustal tectonothermal regime. The volcanic rocks vary from basalt and trachyandesite to trachyliparite from 180 to 140 Ma, showing that the sources of magmas and the thermal levels of partial melting became shallower. With decrease in the crusts’ permeability, the degree of crustal contamination became stronger. The volcanic rocks changed from having no Eu anomalies to having negative Eu anomalies, indicating a crustal thinning process. The Gottini index (t) and differentiation index (DI) of the Early Cretaceous volcanic rocks decreased (Table 1), indicating that the permeability of thickened and folded crust and the geothermal activity decreased. This implies that the depth of the magma source became greater and that crust – mantle interactions became stronger. Therefore, intermediate to basic alkalic volcanic rocks formed at about 120 Ma. However, compared with the basalt of the Early Jurassic Nandaling Formation, the trachybasalt of the Early Cretaceous Donglanggou Formation is on a small scale, and volcanism was weak. The rocks have higher K2O and K2O þ Na2O contents and higher magma differentiation indexes than the Jurassic basalts. The K2O and Na2O þ K2O contents changed from 1.89 to 2.76% and from 3.74 to 3.46%, respectively. The light rare earth elements become more enriched and negative Eu anomalies become more obvious. All these features indicate a crustal extension process, although the crust did not change to the same state as that in the stage of early rifting. However, crustal maturity increased. In addition, the strength of the lower crust decreased with increasing temperature at the crust –mantle boundary, and the lower crustal materials were laterally displaced. The Moho became flatter. Some relicts were carried, absorbed and/or re-formed by the upwelling basaltic magmas, and then carried into the asthenospheric mantle by the locally convecting magma (Nelson 1991). This is the process of lithospheric thinning. The evolution of the Mesozoic Yanshan orogen provides important information on the evolution of intracontinental orogens. They form in the upper crust as a response to strong disturbances of the lithosphere at different levels. The tectonic process consists of thickening and subsequent thinning of the crust and lithosphere, including not only the process of rock deformation but also the building of a mountain belt. Intracontinental orogenies are not restricted to deformation along plate boundaries, but because the driving forces are from the mantle, they may form within previously stable plates. This work was partially funded by the National Natural Science Foundation of China (grant 40372103).

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Upper crustal response to Mesozoic tectonism in western Liaoning, North China, and implications for lithospheric delamination T. D. COPE1 & S. A. GRAHAM2 1

Department of Geosciences, DePauw University, Greencastle, IN 46135, USA (e-mail: [email protected]) 2

School of Earth Sciences, Stanford University, Stanford, CA 94305, USA

Abstract: A widespread and well-documented episode of Late Jurassic– Early Cretaceous rifting followed multiple events of mid- to late Mesozoic crustal contraction in NE China. This extensional deformation was closely associated with widespread Mesozoic magmatism, thought to be related to lithospheric delamination and destabilization of the previously stable North China craton. Early Cretaceous rift-related sedimentary basins in the western Liaoning region of NE China comprise numerous discrete, largely lacustrine half-graben basins bounded by NWrooting low-angle normal faults that sole into older thrusts or mid-crustal shear zones. These basins characteristically lack post-rift thermal subsidence and significantly postdate most of the Mesozoic volcanism in the region. Instead, magmatism that has been attributed to lower crustal foundering, and hence lithospheric delamination (perhaps as old as 160 Ma) accompanied continuing crustal thickening in eastern North China. Thus, although widespread magmatism plausibly played a role in thermally weakening the crust prior to extension, there is little upper crustal evidence that wholesale removal of the lithosphere and lower crust occurred during Mesozoic time. The expansive Cenozoic rift basins of Eastern China, which do contain thick post-rift sequences, constitute a more viable response to lithospheric delamination.

The North China block is an Archaean crustal fragment of Gondwanan affinity that became incorporated into Eastern Asia during Mesozoic time (Zhang et al. 1984; Yin & Nie 1996). Commonly referred to as the North China ‘craton’, the eastern half of the North China block lacks the long-term stability that defines cratons worldwide. Specifically, regions east of the stable Ordos plateau (O, Fig. 1) experienced widespread magmatism and tectonism beginning in Triassic time that was probably related to one or more plate collisions along its northern and southern margins, coupled with palaeoPacific subduction along its eastern edge (Yin & Nie 1996; Davis et al. 1998, 2001; Davis 2003). Geophysical, geochemical, and petrographic evidence suggests that the lower crust and lithosphere were delaminated throughout a vast area of North China, possibly during Mesozoic time (Chen et al. 2003; Gao et al. 2004). Kimberlite-hosted xenoliths of Ordovician age from the eastern part of the North China block indicate the existence of thick (c. 180– 200 km) lithosphere beneath North China until the early Paleozoic (Fan et al. 2000). Basalt-hosted xenoliths of Cenozoic age indicate a thin (c. 80 km) lithosphere underlain by depleted asthenospheric mantle (Zheng et al. 2005), indicating that the lithosphere in Eastern China was significantly thinned prior to Cenozoic time. Gao et al. (2004) suggested that lithospheric foundering occurred during the late Mesozoic, as indicated by the

presence of 159 Ma adakitic rocks that have geochemical and petrological signatures consistent with derivation by partial melting of Archaean lower crustal eclogite. The purpose of this paper is to place upper crustal geological constraints on the dynamics of lithosphere removal in North China. Late Jurassic magmatism considered to be related to lithospheric removal has been recognized in the eastern Yanshan fold– thrust belt, located in western Liaoning Province of NE China (Gao et al. 2004). This region experienced multiple episodes of orogenesis during early Mesozoic time, which was followed by widespread extension in the Early Cretaceous. Both the early Mesozoic orogenesis and the extension that followed were intimately associated with widespread 180– 110 Ma volcanism and plutonism (Davis et al. 2001; Davis 2003).

Tectonic setting Two major crustal elements of NE China form the basement for widespread rifting in the Cretaceous. The Archaean North China block extends northward to the Suolon (or Suolonker) suture, which separates it from the ‘Altaid’ Paleozoic arc terranes and microcontinental fragments that underlie much of Manchuria and Mongolia (Zhang et al. 1984; Wang & Mo 1995; Sengo¨r & Natal’in 1996;

From: ZHAI , M.-G., WINDLEY , B. F., KUSKY , T. M. & MENG , Q. R. (eds) Mesozoic Sub-Continental Lithospheric Thinning Under Eastern Asia. Geological Society, London, Special Publications, 280, 201–222. DOI: 10.1144/SP280.10 0305-8719/07/$15 # The Geological Society of London 2007.

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Fig 1. Generalized map of NE China –eastern Mongolia, showing: (1) major Mesozoic-Cenozoic extensional basins (stippled pattern): E, Erlian basin; G, Gobi basin; H, Hailar basin; NC, North China basin; S, Songliao basin; (2) the stable Ordos block (O, lined pattern); (3) locations of major intraplate thrust belts (barbed lines); (4) Cretaceous metamorphic core complexes (MCCs): Y, Yunmeng Shan MCC; L, Liaonan MCC; W, Waziyu MCC; H, Hohhot MCC; O, Yagan–Onch Hayrhran MCC. Location of Figure 2 is shown by the box.

Yin & Nie 1996; Lamb & Badarch 1997; Cope et al. 2005). Amalgamation of these blocks was completed by Permian–Triassic time (Yin & Nie 1996; Xiao et al. 2003), forming a unified North China plate. Extensive volcanism and plutonism characterized outboard regions of the North China block throughout most of Mesozoic time, and may have been related to palaeo-Pacific subduction (Xu 1990; Davis 2003). Continuing continental contraction along the northern and southern margins of the North China plate, coupled with subduction along its eastern margin, resulted in widespread Mesozoic intracontinental orogenesis (Davis et al. 1996; Yin & Nie 1996; Zhang et al. 1996; Zheng et al. 1998; Darby et al. 2001). The Yanshan fold–thrust belt is an exposed fragment of the orogenic system that developed in response to these events (Fig. 1). The region experienced at least two episodes of contractile tectonism that are manifested by regional Jurassic unconformities (at c. 180 and 160 Ma) and cross-cutting relationships (Davis et al. 2001). The orogenic fabric of the belt is overprinted by widespread midCretaceous extension, including numerous metamorphic core complexes (Fig. 2; Davis et al. 1996, 2002; Webb et al. 1999; Darby et al. 2004; Liu et al. 2005) and their low-strain equivalents (Li et al. 1984, 1995; Ren et al. 2002; Meng et al. 2003). Many of the basins formed during this

episode were later inverted, possibly as a result of mid-Cretaceous or later transpression (Tian et al. 1992; Graham et al. 2001). All of these events are recognized in the Liaoxi (western Liaoning) region of NE China. The following discussion outlines evidence and age constraints for these events in Liaoxi (Fig. 3), with special attention paid to the structural style and facies associated with rift basin development. In Liaoxi, at least two major episodes of thrusting are evident. The first of these involved formation of a SE-vergent brittle nappe, the overturned limb of which is exposed through anticlinal windows in Jurassic volcanic strata (173 Ma, Fig. 3; Davis et al. 2001), which unconformably overlie the structure. This nappe, and the overlying Jurassic volcanic rocks, are cut by a number of younger thrust faults, all of which trend NE –SW and dip to the NW (Fig. 3; Davis et al. 2001; Wang et al. 2001).

Pre-Late Jurassic tectonism and sedimentation Structures and basins associated with a major early Mesozoic tectonic event are exposed throughout the Yanshan region. Sedimentary rocks of Triassic – Early Jurassic age within the Yanshan region are typically boulder to cobble conglomerate (although finer-grained sandstone equivalents do occur), and crop out along the axis of the Yanshan belt (Fig. 3). In the Liaoxi region, two important exposures of these strata and the structures that bound them occur. Northwest of Jianchang (J, Fig. 3), overturned Triassic sandstone is overridden by overturned Cambrian and Ordovician carbonate rocks in the upper plate of a low-angle thrust fault. These upper plate rocks are interpreted as the lower, overturned limb of a brittle, SE-directed recumbent nappe that deforms Triassic and older rocks over a broad area in western Liaoning Province. The overturned limb and floor thrust of the nappe are exposed in a series of erosional windows through unconformably overlying Middle Jurassic and Cretaceous strata. Middle Jurassic volcanic rocks yielding a plagioclase 40Ar/39Ar isochron age of 173 + 4 Ma unconformably overlie both of these units and the thrust fault that separates them, placing an upper limit on the age of faulting (Davis et al. 2001). Locally, Triassic strata in the footwall of the nappe are juxtaposed against the overlying Jurassic strata by younger SE-vergent reverse faulting (Fig. 3). Overturned Cambrian – Ordovician strata associated with the nappe are exposed intermittently along strike before plunging into the subsurface west of Dachengzi (D, Fig. 3). Southwest of Lingyuan (L, Fig. 3), syntectonic strata record folding and thrusting events that are

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Fig 2. Generalized tectonic map of the study area, showing thrust trends in the Yanshan fold– thrust belt, major Cretaceous sedimentary basins and bounding structures, and Cretaceous volcanic strata. Arrows indicate Cretaceous extension directions derived from stretching lineation in the Yunmeng Shan metamorphic core complex (YMCC; Davis et al. 1996), fault striae on the Louzidian normal fault (LF; Han et al. 2001), and Liaonan MCC (LMCC; Yin & Nie 1996). Cretaceous sedimentary basins: CF, Chifeng basin; CY, Chaoyang basin; F, Fuxin basin; J, Jianchang basin; L, Luanping basin; JF, Jiufotang basin. Also shown: WMCC ¼ Waiziyu MCC (Darby et al. 2004); C ¼ Chengde. Stars indicate radiometric ages (in Ma). †Swisher et al. (1999); *Davis et al. (1998); **Niu et al. (2003). Location of Figure 3 is shown.

broadly time-correlative with formation of the recumbent nappe described above. Proterozoic – Ordovician carbonate strata in this region are deformed into broad, NE-plunging folds that are unconformably overlain by coal-bearing cobble conglomerate designated Middle Jurasssic in age (T?, Fig. 3; Liaoning Bureau of Geology and Mineral Resources (LBGMR) 1989). Immediately to the west of these strata, a vertical to steeply SE-dipping fault places highly deformed basal Proterozoic quartzite above an eastward-thickening, synformally folded section of carbonate-clast cobble to boulder conglomerate (T, Fig. 3; Davis et al. 2001). The Proterozoic rocks in the hanging wall of this fault are pervasively deformed into tight, NW-overturned folds, suggesting that this structure is a NW-directed thrust. The conglomerate that lies in the footwall is up to 2400 m thick, extremely coarse, and contains an unroofing sequence

reflecting progressive denudation of Ordovician, Cambrian, and Proterozoic source strata (Sun 2002). In addition, an intact block of Cambrian – Ordovician carbonate strata that lies in the centre of the basin has been interpreted as a massive, 5 km wide slide block (Sun 2002), further supporting the notion that these strata were deposited during large-scale tectonism in the Yanshan region. Biotite from a dacite that conformably underlies this sequence yields an 40Ar/39Ar plateau age of 219.4 + 1.0 Ma (Fig. 4a), placing a lower limit on the age of this unit. Although the structures that controlled sedimentation into this basin have not been identified with certainty, these strata clearly record an early phase of intense tectonism and denudation in the Yanshan belt. This phase of deformation in the Yanshan was probably related to closure of an ocean basin and amalgamation of North China with volcanic arc

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Fig 3. Geological map of western Liaoning (Liaoxi), showing major rock units and radiometric ages discussed in the text. Unlabelled or queried rock units have not been dated; ages for these units are from LBGMR (1989). City names: B, Beipiao; C, Chaoyang; D, Dachengzi; F, Fuxin; J, Jianchang; L, Lingyuan; Y, Yixian. References for age determinations are as follows: 1Davis et al. (2001); 2Gao et al. (2004); 3Swisher et al. (1999, 2002); 4Chen et al. (1997); 5Wang et al. (2001); 6Darby et al. (2004). Locations of seismic sections in Figures 6, 7 and 12 are shown. Locations of measured stratigraphic sections denoted by palaeocurrent localities are indicated with an open asterisk. N, number of measurements.

terranes in Mongolia. The resulting suture, variously termed the Suolon suture (Wang & Mo 1995), Solonker suture (Ruzhentsev et al. 1985), Junggar–Hegen suture (Zhang et al. 1984), or Tian Shan– Yin Shan suture (Yin & Nie 1996), separates Archaean crust of the North China block from the late Paleozoic ‘Altaid’ arc terranes to the north (Sengo¨r & Natal’in 1996; Xiao et al. 2003). Permian plutonic rocks south of the suture zone are inferred to be arc plutons related to a southwarddipping pre-collisional subduction zone (Wang & Liu 1986; Cui & Wu 1997; Cope et al. 2005).

Late Jurassic – Early Cretaceous tectonism and sedimentation A regional, pre-173 Ma surface overlaps all older thrust structures and associated basins in the Yanshan belt (Davis et al. 2001). The strata that

overlie this unconformable surface consist of andesitic –rhyolitic volcanic rocks that yield ages between 173 and 155 Ma (Fig. 3) and are cut by a strongly SE-vergent system of thrust faults. Northwest of Jianchang (J, Fig. 3), the overturned limb of the brittle nappe discussed in the preceding section, and the volcanic strata that overlie it, are cut by one or more SE-vergent thrusts that place Proterozoic strata above volcanic rocks as young as 156 Ma (Fig. 5a; Davis et al. 2001). Southwest of Lingyuan, the Triassic –Early Jurassic basin discussed above, and the NW-directed thrust that bounds it, have been truncated by a younger SE-vergent thrust that places these units above Middle Jurassic conglomerate to the SE (Fig. 3). Southeast of Jianchang, another south-directed thrust fault places Proterozoic carbonate above Jurassic volcanic rocks that may be correlative with the 173–155 Ma volcanic rocks to the NW. Jurassic volcanic units are overlain by alluvial, fluvial and aeolian strata assigned to the Late

UPPER CRUSTAL RESPONSE TO MESOZOIC TECTONISM (a)

(b)

(c)

Fig 4. 40Ar/39Ar ages discussed in the text. (a) Sample 01LY160, biotite from a dacite flow that underlies the base of the foredeep sequence beneath the NW-vergent thrust fault south of Lingyuan. (b) Sample 01DZ112, biotite from a reworked tuff above the overturned limb of the SE-vergent nappe south of Dachengzi. (c) Sample 00LP606, biotite from the welded tuff that underlies the Cretaceous fill of the Luanping basin. Errors bars are 1s.

Jurassic Tuchengzi Formation (LBGMR 1989). Because Late Jurassic–Early Cretaceous sedimentation within the Yanshan region is markedly diachronous, the usage of this designation in this paper is

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restricted to those strata that are continuous with dated strata in the Liaoxi region. At least the upper part of these deposits are Early Cretaceous in age: a tuff interbedded with the upper Tuchengzi Formation (J3–K1, Fig. 3) yields an 40Ar/39Ar age of 139.4 + 0.1 Ma (Figs 2 and 3; Swisher et al. 2002). Tuchengzi strata in the Liaoxi region were deposited primarily within a single, stratigraphically continuous basin that spans the central portion of Figure 3. Strata within this basin are primarily flat-lying fine-grained clastic rocks that thicken towards the leading edge of the thrust belt and are upturned against it in a manner strongly suggestive of a flexural foreland basin setting (Fig. 6). The thickening shown on the seismic line in Figure 6 is not accompanied by coarsening, however: these rocks are the finest-grained Mesozoic strata in the Yanshan region. For example, Late Jurassic–Early Cretaceous strata that unconformably overlie deformed Proterozoic– Paleozoic rocks NE of Jianchang (J, Fig. 3) are composed chiefly of shale with only thin lenses of volcanicderived pebble conglomerate and sandstone (Fig. 5b). The single most distinctive unit within the basin is a vast sheet of cross-stratified green aeolian sandstone that caps the exposed part of the Jurassic –Cretaceous section and constitutes the major lithology in the region, even where proximal to the exposed basin margin (Fig. 5c). If these strata were deposited during a phase of active tectonism in the Yanshan belt, they must have occupied a distal position relative to uplifted regions. Tuchengzi strata are cut by thrust faults only along the northwestern margin of the basin (Fig. 3), and here also are relatively fine-grained sandstone and pebble conglomerate, suggesting that thrusting in this region operated late in the history of the Yanshan belt and was perhaps unrelated to the more widespread SE-vergent thrusts discussed above. It is unclear what drove this event, which is time-equivalent with structures throughout northern North China (Davis et al. 1998, 2002; Zheng et al. 1998; Darby et al. 2001). It may be a late effect of North China–South China collision, palaeo-Pacific subduction, or collision of the amalgamated North China plate with Siberia (Davis et al. 2001).

Mid-Cretaceous extension Late Jurassic–Early Cretaceous contractile deformation in the Yanshan belt was followed by widespread and well-documented mid-Cretaceous extension throughout the northern part of the North China block and surrounding areas (Webb et al. 1999; Graham et al. 2001; Davis et al. 2002; Ren et al. 2002; Meng 2003; Meng et al. 2003). In the Yanshan belt, variably deformed

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Fig 5. Outcrop photographs of Jurassic deformation and strata in the Liaoxi region. (a) Recumbent brittle nappe developed in Cambrian –Ordovician carbonate rocks (C-O). View to the SW. The floor thrust places these strata above overturned Triassic rocks (Tr), which are thrust upon Middle Jurassic strata (J2). The nappe is cut by a younger thrust, which places Proterozoic strata (PC) above its overturned limb. Field of view c. 5 km. (b) Fine-grained Late Jurassic– Early Cretaceous strata exposed c. 50 km NW of Jianchang. View to the east. Beds dip away from the viewer at c. 608. The base of the section consists largely of shale; the uppermost beds are fine- to medium-grained sandstone. (c) Distinctive green, aeolian, fine- to medium-grained sandstone unit within the Tuchengzi Formation. (Note the 4 m high, steeply dipping cross-stratification within the central bed in the photograph.)

Fig 6. NW–SE time-migrated seismic profile (see Fig. 3 for location) through Tuchengzi Formation units east of Chaoyang. (Note substantial thickening of Jurassic– Cretaceous units towards the west, where they are upturned against the thrust belt.) TWTT, two-way travel time.

Late Jurassic–Early Cretaceous and older rocks are overlain with profound angular unconformity by relatively undisturbed mid-Cretaceous andesitic – rhyolitic volcanic, volcaniclastic, and lacustrine strata (Davis et al. 1998, 2001). South of Chengde, in Hebei Province (C, Fig. 2), volcanic rocks yielding an age of 135 Ma unconformably overlie synformally folded Late Jurassic strata (Niu et al. 2003). In Luanping basin (L, Fig. 2), a rhyolite tuff that disconformably overlies Late Jurassic strata yields an 40Ar/39Ar biotite age of 130.8 +0.5 Ma (Fig. 4c). In the Liaoxi region, Early Cretaceous volcanic and lacustrine strata of the Yixian Formation unconformably overlie Late Jurassic – Early Cretaceous Tuchengzi Formation strata and yield an 40Ar/39Ar sanidine age of 124.6 +0.1 Ma (Figs 2 and 3; Swisher et al. 1999). Volcanic rocks of Early Cretaceous age underlie lacustrine

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strata that reside within numerous small half-graben basins that developed within the Yanshan belt and along its periphery. The style of mid-Cretaceous extension in NE China encompasses two end-member types: (1) low-strain extension, which resulted in the formation of numerous small half-graben basins throughout the Yanshan region, NE China, and Mongolia; (2) high-strain extension, which resulted in the formation of a linear belt of metamorphic core complexes along an east –west zone that roughly parallels the axis of older compression (Davis et al. 1996, 2002; Webb et al. 1999; Darby et al. 2004; Liu et al. 2005).

Low-strain structures and basins The geometry and occurrence of low-strain extensional basins in the Liaoxi region was in part controlled by the geometry of older structures. Figures 7 and 8 show two serial time-migrated seismic profiles, calibrated to surface geology, that transect extensional basins developed atop the Yanshan thrust belt. The northern of these lines (Fig. 7) portrays two NW-dipping listric normal faults that bound half-graben basins containing Lower Cretaceous volcanic and clastic strata. One of these faults (the southeastern of the two), is steeper (c. 458 in surface exposures), and has developed a significant antiformal geometry in its hanging wall that affects pre-rift and synrift strata consistent with rollover anticlines that develop in

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the hanging walls of listric normal faults. Although lack of well control does not permit identification of the Proterozoic –Cambrian contact at depth, this contact has been approximated on the seismic line using known stratigraphic thicknesses, and it ties (with a great deal of uncertainty) to the same contact exposed in the footwall of the fault, yielding extension of c. 4–5 km across the structure. The northwestern portion of the seismic line images an entirely different extensional geometry. Here, a half-graben basin is bounded by a low-angle normal fault that parallels an older thrust in surface exposures. This thrust places Cambrian and Ordovician rocks above red sandstone and shale that resemble Triassic strata exposed to the south in the footwall of the brittle nappe. Welldefined reflectors that are interpreted as the surface of the thrust are not cut by the normal fault, and it is interpreted that the normal fault soles into and reactivates this older thrust at depth. Although stratigraphic separation cannot be determined across this structure, restoring the base Cretaceous to the surface yields a minimum of 4 km extension. Another seismic profile (Fig. 8) displays a similar set of features that are imaged to nearly 4 s two-way travel time (TWTT) depth. Surface geology in this region is complex and includes exposures of the brittle nappe and unconformably overlying Jurassic volcanic rocks discussed in preceding sections. The seismic line images deeper structural levels occupied solely by SE-vergent

Fig 7. Time-migrated seismic profile through rift basins developed atop the Jurassic–Cretaceous Yanshan fold– thrust belt in the northern part of the study area. The seismic interpretation is tied to surface geological relationships (shown). (See Fig. 3 for location.) It should be noted that the western (shallower) basin-bounding normal fault soles into the older thrust at shallow depth; the steeper eastern fault does not.

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Fig 8. Time-migrated seismic profile through rift basins developed atop the Jurassic–Cretaceous Yanshan fold–thrust belt in the southern part of the study area. Both normal faults shown in this line sole into earlier-formed thrust faults. The fault bounding the Jianchang basin is probably oversteepened at this location as a result of intrusion of the younger pluton to the far east of the line. (See text for explanation.) Scale is the same as in Figure 7.

thrust faults that crosscut older structure and are associated with Late Jurassic deformation. Two major thrusts are visible, both of which deform Proterozoic–Paleozoic strata and one of which displays a ramp-flat geometry. The lowermost (easternmost) thrust fault represents the leading edge of the Late Jurassic –Early Cretaceous thrust belt, and in outcrop places Proterozoic strata against the foreland basin fill to the east. The westernmost thrust cuts the overturned limb, footwall strata, and floor thrust of the Late Triassic –Early Jurassic nappe, schematically shown near the top of the seismic line. On the surface, these thrust faults closely parallel Cretaceous normal faults that bound two clearly imaged half-graben (the Jiufotang and Jianchang basins) in the upper portion of the seismic line. These normal faults do not cut reflectors beneath the older thrust faults that parallel them, and it is inferred that the normal faults sole into the thrusts at depth. Minimum extension, based upon restoration of the base Cretaceous unconformity to the surface, is on the order of 11 km, similar to the amount derived for the basins in Figure 7. The geometry of the two seismic profiles, although similar, differs in terms of fault spacing. Three of the four normal faults imaged in these lines appear to reactivate older thrust structures, and it is therefore likely that the primary control on normal fault spacing is the spacing of these older thrusts. Widely separated normal fault

terminations must be linked by some form of transfer structure between the two lines. Lower Cretaceous volcanic rocks in the northern part of the Jianchang basin are juxtaposed against Proterozoic –Paleozoic carbonate rocks along a steeply dipping, NW –SE-trending structure that is interpreted as a left-lateral transfer fault linking the two basin systems (Fig. 3). The seismic interpretations described above are verified by surface exposures of basin-bounding structures. The normal fault that bounds the southeastern margin of the Jianchang basin is localized along the trace of a gently NW-dipping thrust fault that places Proterozoic carbonate rocks above Jurassic volcanic rocks to the south of the basin (Fig. 3). Along the seismic profile, the normal fault dips 51–548 to the NE and exhibits numerous kinematic indicators showing nearly pure dip slip (Fig. 9a). This segment may have been steepened by 113 Ma intrusion of a pluton into its footwall (Fig. 3). The normal fault flattens to dips of 20 –228 in the southern part of the basin, where it approaches the adjacent thrust and becomes parallel with it (Fig. 9b). The Jiufotang basin is bounded to the SE by a gently NW-dipping normal fault, previously interpreted as an unconformity (LBGMR 1989; Davis et al. 2001). The fault is expressed in outcrop by a 1–3 m wide zone of carbonate fault breccia and gouge along which excellent exposures of the fault surface exist. Striae on

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Fig 9. Outcrop photographs of structural relations along basin-bounding normal faults. (a) Jianchang basin-bounding normal fault. View to NE. The fault places Cretaceous conglomerate (K1cong) against Proterozoic carbonate strata (pC) in the upper plate of a thrust fault that crops out to the east of the photographs. The fault zone contains striae and subsidiary faults indicating a dip-slip sense of motion. A leucocratic intrusive rock (Kgr) parallels the fault zone, and clasts from this intrusive rock, and the surrounding carbonate, are present within the basin-margin conglomerates. (b) Jianchang basin-bounding fault in the southern portion of the Jianchang basin. The fault dips parallel to the older thrust fault. Jv, Jurassic volcanics; K1j, Jiufotang Formation. Jiufotang facies are not exposed along the fault at this locality. (c) Surface relationships between basin-bounding normal fault and Jurassic thrust fault, Jiufotang basin. View to NE. Prominent flatirons in the foreground are the exhumed surface of the basin-bounding normal fault, which overlies and parallels a Jurassic thrust beneath it. The thrust places Proterozoic carbonate strata (pC) above Jurassic volcanic rocks (Jv), which unconformably overlie Paleozoic carbonate rocks (C-O) in the upper plate of the brittle nappe discussed in the text. The two faults merge at the bottoms of some of the deeper gullies.

this surface indicate nearly pure dip slip. The surface trace very closely follows the trace of a Jurassic thrust fault that places Proterozoic carbonate rocks above Jurassic volcanic rocks that unconformably overlie the upper and lower plates of the Triassic nappe lying to the east (Fig. 9c). Over most of its length, the plane of the fault dips 20– 368 NW, nearly parallel to the adjacent thrust. In many places the upper plate of this older fault is reduced to a sliver by normal faulting, and locally it is cut out entirely; the maximum distance between the two faults at the surface is c. 500 m. The Jiufotang normal fault can be traced in outcrop to the NE for several tens of kilometres before its trace becomes lost under Quaternary alluvium (Fig. 3).

Basin-fill facies To document the synextensional nature of the fills of half-graben basins in the Liaoxi region, the following discussion outlines facies relations, provenance, and paleodispersal patterns within two well-exposed examples: the Jianchang basin and the Jiufotang basin (Fig. 8). Jianchang basin. The fault-bounded margin of the Jianchang half-graben is flanked by poorly organized cobble to boulder conglomerate that passes laterally away from the fault into finer-grained marginal lacustrine strata. Clast compositions in exposed levels of this basin-margin conglomerate

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directly reflect footwall lithologies opposite the basin-bounding normal fault. Because of the low dip of the strata and lack of significant topographic relief, upsection compositional trends within the conglomerate were not observed. Directly east of Jianchang (J, Fig. 3), the normal fault juxtaposes Proterozoic carbonate in the upper plate of the Jurassic thrust against proximal conglomerate facies in the Jianchang basin. The upper plate of the thrust in this locality is intruded by a 113 Ma granodiorite pluton (Wang et al. 2001) and one or more highly sheared, leucocratic, K-feldspar porphyritic dykes. Conglomerate clast compositions in the Jianchang basin opposite this assemblage consist of Proterozoic carbonate and porphyritic volcanic clasts that texturally and compositionally match the K-feldspar porphyry found in the footwall of the fault. Granodiorite clasts from the pluton are absent, suggesting that this pluton is either younger than the basin fill or was not yet unroofed during deposition of the conglomerate. Further south, where the footwall of the normal fault exposes Jurassic volcanic rocks in the lower plate of the thrust, clast compositions in proximal conglomerate consist entirely of well-rounded andesitic volcanic rocks derived from these lower plate volcanic rocks. This relationship suggests that: (1) some thrust segments were reactivated as normal faults throughout their entire updip extent, exposing lower plate rocks to erosion immediately following inversion; (2) the normal fault locally cut outboard of the frontal trace of the thrust; or (3) upper plate rocks were eroded in the footwall of the normal fault prior to deposition of conglomerate in this locality. The complete lack of carbonate clasts derived from upper plate sources in this conglomerate suggests that the first two alternatives are the most likely. Paleocurrents in SE basin-margin conglomerate (largely imbrication) are directed to the NE, normal to the strike of the normal fault (Fig. 3). Basin-margin conglomeratic facies extend to c. 1 km from the fault, where they interfinger with coal-bearing marginal lacustrine rocks. Coal is mined along a discontinuous linear trend at the junction between basin-margin conglomeratic facies and predominantly finer-grained lacustrine facies exposed c. 1 km basinward from the fault. Coal probably occupied interdistributary mires between fan-delta complexes feeding into the lacustrine system from the faulted margin (see Whateley & Jordan 1989). Lacustrine facies associations, well exposed in the basin axis at Jianchang, comprise siltstone and shale interbedded with pebble conglomerate and coarse- to medium-grained sandstone (Fig. 10). These interbedded units are punctuated by the conglomeratic deposits of multiple coarse-grained Gilbert-type deltas up to 12 m thick, which form the tops of crudely developed

coarsening-upward successions (Fig. 11a). The 130 m of exposed strata in this area display an upward shallowing trend in which deltaic units increase in frequency at the expense of openlacustrine shale and siltstone. Five lithofacies are recognized in the measured section (Fig. 10) and are codified using a modified nomenclature after Miall (1978). (1) Grey–green lacustrine shale and siltstone (Fl), which are interbedded with thin sandstone and conglomerate units, constitute the dominant lithology. These finegrained intervals are thinly bedded to laminated, commonly bioturbated, and contain abundant woody debris, leaf impressions, and rare bivalve fossils and fish scales. Laminated shale and siltstone containing aquatic faunal remains are indicative of open-lacustrine conditions in which sedimentation took place through settling of clay or silt from suspension (e.g. Talbot & Allen 1996). (2) Sandstone units (Sm) interbedded with this lacustrine shale and siltstone are thinly bedded (commonly ,10 cm), coarse- to medium-grained, and increase in frequency and coarseness upwards. Sandstone beds are commonly current rippled, plane laminated, and/or graded where bioturbation has not destroyed primary sedimentary structures. This facies is interpreted as sublacustrine turbidite deposits (e.g. Bouma 1962). (3) Massive, matrixsupported pebble conglomerate beds 10–150 cm thick (Gmm) occur within shale- and siltstonedominated successions. These are clearly the products of cohesive debris flows (e.g. Lowe 1982). (4) Clast-supported conglomerate (Gcm) occurs interbedded with lacustrine shale as well. Conspicuous features of these clast-supported conglomerate beds include a complete lack of internal traction structuring or bedding, poor internal organization, and lack of basal scour or intraformational rip-up clasts indicative of erosional power. Nearly all of the conglomerate units observed overlie shale or siltstone along a flat, non-erosive contact that parallels bedding in the underlying strata. These conglomerate units are interpreted as the products of sublacustrine clast-rich debris flows (e.g. Nemec et al. 1984). (5) Planar cross-stratified conglomerate beds 6–12 m thick (Gp) cap a number of weakly developed coarsening-upward successions and contain spectacular, high-angle (25 –308), nontangential foresets composed of alternating sand – gravel couplets that penetrate the entire bed (Fig. 11a). The bases of these units are non-erosive. Some foreset beds are matrix supported, but most are clast supported and display imbrication oriented down the foreset dip. Individual sand- and graveldominated foresets are internally inverse graded, and coarser clasts are concentrated at the toes of foresets, as is typical of foresets generated by avalanching grain flows (Nemec 1990). These units

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Fig 10. Stratigraphic column, Jianchang basin. Lithofacies abbreviations, modified from Miall (1978), are discussed in the text. Paleocurrent arrows for facies Gp, interpreted as Gilbert-type delta deposits, are shown to the right of the graphic column.

are interpreted as the foresets of coarse-grained Gilbert-type deltas. Gilbert delta complexes are typically underlain by a package of bottomset strata 5–10 m thick composed of plane-laminated coarse sandstone interbedded with matrix-supported conglomerate, interpreted as the products of proximal sediment gravity flows emanating from the toes of the prograding delta. Topset strata typically consist only of a thin pebble to cobble lag that overlies the tops of the foresets and is draped by lacustrine shale representing the base of the next sequence. Northeast-directed palaeocurrents, as derived from foreset dip, indicate that these Gilbert delta complexes prograded axially across the basin from the SW, subparallel to the basin margin. The maximum thickness of these units indicates a water depth of less than 12 m during delta progradation in this part of the basin. The normal fault that bounds the Jianchang basin is linked, along a left lateral transfer structure, to another normal fault that continues northward to

the latitude of Chaoyang (C, Fig. 3). Facies observed in reconnaissance adjacent to this normal fault reveal a similar assemblage of alluvial cobble to boulder conglomerates, derived from footwall lithologies, which clearly reflect the syntectonic nature of the basin fill. Along this fault, Cretaceous strata consist of conglomerate (with clasts up to 90 cm in diameter) that is interbedded with laminated to thinly bedded siltstone and shale. Clast lithologies in these units include dolomite, chert, and limestone derived from Proterozoic and lower Paleozoic units in the footwall of the fault. Northeast of Chaoyang (C, Fig. 3), Cretaceous boulder conglomerate yielding NW-directed palaeocurrents is juxtaposed against deformed carbonates along an east –west- to NE– SW-striking normal fault, which dips c. 408 NW. Carbonate rocks in the footwall of this fault reside in the upper plate of a system of thrust faults that cut Jurassic strata to the south. Cretaceous conglomerate north of the normal fault consists entirely of clasts

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Fig 11. Outcrop photographs of basin-fill facies in the Jianchang and Jiufotang basins. (a) Gilbert delta foresets, Jianchang basin. Brick wall is 2 m high. Foreset beds are underlain and overlain by shallow lacustrine strata. View to NW. (b) Basin margin alluvial conglomerate, Jiufotang basin. The circled clast of Proterozoic dolomite is 1.5 m in diameter. (c) Slump or debris-flow bed, Shanzuizi basin. Well-bedded units in the upper right of the photograph are rafts of coherent strata near the top of the debris-flow unit. (d) Late Cretacous thrust fault south of Beipiao. View to north. Person (arrowed) for scale. The thrust places Proterozoic dolomite (pC) above Upper(?) Cretaceous conglomerate and sandstone (K2?).

of green sandstone, resedimented pebble conglomerate, and andesitic volcanic rocks that are remarkably similar to Jurassic units found in the foredeep of the Jurassic thrust belt (J3–K1 and Jv, Fig. 3). Clasts of Proterozoic and Paleozoic lithologies (which form the now exposed footwall of the adjacent normal fault) are absent, suggesting that these units were once overlain by Jurassic strata that served as a source for conglomerate during Cretaceous time. Jiufotang basin. The Jiufotang half-graben lies to the NW of the Jianchang half-graben, and contains c. 1 km of sedimentary fill (Fig. 8). The faulted margin of the Jiufotang half-graben is dominated throughout its length by a poorly organized, angular, cobble to boulder conglomerate (Fig. 11b). The most poorly sorted conglomerate occurs nearest the fault zone, where bedding is indistinct, and clasts are mostly angular and up to 2.5 m in length occur. This facies is interpreted as subaerial talus slope deposits. Clast compositions in basin margin conglomerate comprise a mixture of carbonate clasts (derived from the upper plate of the footwall

thrust) and volcanic clasts (derived from the lower plate), with rare sandstone and granitic clasts. Both the granite and sandstone clasts were probably recycled from Jurassic conglomerate and sandstone overlying the volcanic sequence in the lower plate of the Jurassic thrust. Three stratigraphic columns were measured 1–2 km from the basin-bounding normal fault and spaced c. 500 m apart (Fig. 12). The facies described in these columns consist of lacustrine shale interbedded with pebble to boulder conglomerate and coarse- to medium-grained sandstone, interpreted as the proximal and distal deposits of one or more basin-margin fan-delta complexes, respectively. The three columns document progressive fining and increasing lacustrine influence in a basinward direction from the basin-bounding normal fault, which lies to the SE. Minor structural complexities, rapid facies changes, and pervasive synsedimentary deformation rendered accurate correlation of these three columns problematic, although all three lie within in the same level of exposure in generally flat-lying, unfaulted rocks. Seven main lithofacies are recognized in the measured sections (Fig. 12). (1) Laminated to

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(a)

(b)

(c)

Fig 12. Measured sections, Jiufotang basin margin. (Note horizontal and vertical scale.) The fault-bounded basin margin is c. 500 m SE of section (c). (See text for discussion.)

thinly bedded shale and siltstone (Fl) represent lacustrine settling of fines from suspension (Talbot & Allen 1996). Wood fragments, delicate leaf

imprints, and fish scales are common within this facies. Approximately 1 km NW of the measured section, beds of siltstone are deformed into

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small-scale, NW-overturned slump folds that reflect continued influence of a NW-facing slope. (2) Massive, graded, plane-laminated, and/or rippled sandstone (Sm) commonly occurs interbedded with lacustrine shale and siltstone. Beds are laterally continuous and 5– 50 cm thick. A majority of sandstone beds of this facies are clearly turbiditic and contain Bouma divisions Ta, Tb and Tc. Pebbly and granular bases to graded beds are common. Grain size varies from very coarse to medium sand. (3) Massive, clast supported conglomerate (Gcm) occurs in beds 0.5–10 m thick with non-erosive, but commonly loaded, bases. These units are very poorly sorted with highly angular clasts up to 50 cm in diameter. Some units exhibit weakly developed imbrication in clasts with long axes parallel to flow direction, indicative of clast transport in suspension (Harms et al. 1975). These units are interpreted as cohesionless, clast-rich debris flows or subaqueous grain flows (Lowe 1982; Nemec & Steel 1984). (4) Mud matrix-supported conglomerate units (Gmm), interpreted as cohesive debris-flow deposits, commonly interfinger with chaotic slumped intervals (Fig. 11c). Individual matrix-supported conglomerate beds attain thicknesses of up to 8 m, have non-erosive bases (although basal load features are common), and contain very large, deformed and broken blocks of intraformational sandstone, conglomerate, and shale clearly derived from slope failures up depositional dip. Larger clasts and blocks are typically concentrated near the tops of beds as rafts, and commonly project from the upper surface of the bed into the overlying strata, a common trait of cohesive debris-flow deposits (Nemec & Steel 1984). Weakly developed inverse grading is common. The bases of some beds contain thin, clast supported and inversely graded layers interpreted as basal traction carpets (e.g. Lowe 1982). (5) Large-scale planar cross-stratified conglomerate (Gp) occurs only in measured section B. It consists of interstratified pebble to cobble conglomerate and very coarse sandstone. Cross-stratification reaches amplitudes of up to 8 m. These units are similar to the Gilbert delta foreset deposits in the Jianchang basin. (6) Imbricated clast-supported conglomerate (Gcmi) consists of well-imbricated cobble to boulder conglomerate interstratified with trough cross-stratified or planelaminated coarse sandstone. Imbricated clasts have long axes transverse to flow, indicating bedload traction transport by rolling (Harms et al. 1975). Bedding style is tabular to broadly lenticular, and typically consists of alternating packages of normally graded to ungraded conglomerate overlain by sandstone 1–2 m thick. This facies is interpreted as subaerial sheetflood deposits (see Blair 1999) deposited on the upper reaches of the fan-delta

complex. (7) Clast-supported, trough crossstratified conglomerate and sandstone (Gct) occur in close association with facies Gcmi. Trough crossstratification is characteristically low-angle, indicating a highly erosive environment. Bedding is discontinuous and internal scour is common. These units are interpreted as alluvial fan channel deposits. The facies documented within the measured sections shown in Figure 12 represent a complex marginal lacustrine environment dominated by mass wasting, and match well with facies documented from other fan-delta settings (Wescott & Ethridge 1980; Postma 1984, 1990). The distribution of facies in each column is dependent on proximity to the basin-margin fault. The coarse-grained subaerial alluvial fan facies of column C, located c. 1 km from this fault, pass basinward into the shaledominated sublacustrine successions of columns A and B. Planar cross-stratified conglomerate, possibly representing the deposits of small, Gilbert-type deltas, occurs only in column B and marks the temporary position of the lacustrine shoreline. Debrisflow conglomerate, dominantly clast-supported in columns B and C, is largely matrix-supported in column A, perhaps reflecting greater runout distances of cohesive debris flows over clast-rich suspensions supported by dispersive pressure, which require a high slope to move (Nemec 1990). Slumped intervals occur within all three columns, reflecting the presence of a failure-prone slope. Paleocurrent indicators, largely derived from imbrication, indicate sediment transport to the NW, away from the basin margin fault. The poorly exposed ramp margin of the Jiufotang half-graben is dominated by lacustrine shale floored by a thin interval of fine-grained meandering fluvial sandstone. These dip gently SE and lie unconformably upon either folded Cambrian – Ordovician carbonate rocks or highly altered Cretaceous volcanic strata, which floor the basin to the NW. Further north, exposures of SE-dipping strata that unconformably overlie Paleozoic carbonates consist of highly channelized, trough crossstratified sandstone and pebble conglomerate interpreted as fluvial deposits. These yield SE-directed palaeocurrents (Fig. 3). They are sparsely exposed north of Dachengzi and west of that locality towards Lingyuan.

High-strain extension Seismic data within Fuxin basin (F, Figs 2 and 3) image a low-angle (c. 208), NW-dipping normal fault that has been variably interpreted as the master detachment of a metamorphic core complex (the Waziyu MCC) of mid-Cretaceous age (Darby et al. 2004), or as a younger, steeper

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normal fault that truncates a sinistral ductile shear zone within the core of the adjacent Yiwulu Shan (Zhang et al. 2003). The overlying Fuxin basin has been interpreted (by Darby et al. 2004) as a supradetachment basin that developed as the result of slip on the NW-rooting Waziyu detachment. The seismic data (Fig. 13) image a gently dipping reflector that separates highly reflective, flat-lying units above from poorly reflective (or non-reflective) units below, in a manner consistent with juxtaposed high and low metamorphic grade. Below the base of the Cretaceous fill of Fuxin basin (Base K1, Fig. 13a), this reflector truncates a series of SE-dipping reflectors before appearing to sole into a set of flat-lying reflectors (possibly artefacts of the seismic processing) at the 4 s (TWTT) base of the seismic line. A strike line

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through the axis of Fuxin basin (Fig. 13b) portrays a curvilinear fault trace that is apparently cut by at least two SW-dipping faults. Because of low reflectivity beneath the Waiziyu detachment, it is unclear whether these structures sole into the master structure, or cut it. Reflectors within the Fuxin basin starkly contrast with those in low-strain basins (Figs 7 and 8) in that they are relatively flat-lying (although highly disrupted) and do not dip into the basinbounding Waziyu detachment. This is a common trait of many of the basins throughout NE China (Meng et al. 2003). These attributes are characteristic of supradetachment basins, where footwall uplift inhibits rollover into the underlying detachment and pushes depocentres outward from the breakaway zone (Friedmann & Burbank 1995).

Fig 13. Seismic lines through the Fuxin basin and bounding normal fault. (a) Dip line. (b) strike line. Notation as in Figure 8. (See text for discussion.)

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Post-extensional inversion Conglomerate, sandstone, and red siltstone mapped as Late Cretaceous in age (LBGMR 1989) crop out locally south of Beipiao (B, Fig. 3). Although the reason for this age assignment is uncertain, these rocks are certainly younger than the uppermost underlying Tuchengzi Formation because they contain abundant, distinctive green sandstone clasts derived from that unit. The contact between the two units is mapped as an unconformity (LBGMR 1989). These Cretaceous strata are overthrust by Proterozoic rocks along a well-exposed south-directed reverse fault that terminates NE of Chaoyang (Fig. 11d). This fault is perhaps the youngest thrust structure in the Yanshan region, particularly if the strata it cuts are indeed Late Cretaceous in age. The hanging wall of this fault is deformed into NNW-plunging, SSW-overturned folds that indicate a south-SW direction of upper plate transport. Along strike and to the west, the thrust merges with a roughly north–south-trending strike-slip fault that offsets both Cretaceous normal faults and the older thrust belt in a left-lateral sense (Fig. 3). No offset equivalent of the Cretaceous thrust exists to the east of this fault, suggesting that this thrust is a transpressional feature related to sinistral strike-slip. It is interesting that this strike-slip fault runs parallel to the Tan-Lu fault to the east (Fig. 1), and to other recognized strike-slip structures of uncertain age in the region with demonstrated left-lateral offset (Fig. 2; Davis et al. 2001; Han et al. 2001). An event of mid- to Late Cretaceous inversion, possibly related to transpression along NE-striking faults, is recognized in many of the basins of NE China (Tian et al. 1992; He & Wang 2004), and in Mongolia (Graham et al. 2001).

Regional timing of extension The facies and geometry of the basins described in this paper match well with those described in other half-graben basins within the Cretaceous extensional province of NE China (Li et al. 1984, 1995; Meng 2003; Meng et al. 2003), which suggests a common tectonic driver. Specifically, they share several common characteristics: (1) humid-climate facies associations, as indicated by the abundance of coal and lack of evaporitic lacustrine facies (Li et al. 1984; Wu et al. 1992); (2) asymmetric, half-graben geometries, which generally lack evidence for extensive post-rift thermal subsidence (Dou 1997; Ren et al. 2002; Meng et al. 2003); (3) flattening of bounding structures at mid-crustal levels, and close association of these structures with older thrust faults (Meng

et al. 2003); (4) NE–SW elongate orientations, subparallel to late Mesozoic inferred continental margin trends (Ren et al. 2002). Together, these basins define a widespread extensional province that developed throughout NE China and SE Mongolia during the late Mesozoic. Over 200 small rift basins developed within the NE China region during the Cretaceous (Dou 1997; Lin et al. 2001). Well-documented examples of late Mesozoic basins related to this rifting event include a number of coal-bearing half-graben basins in Inner Mongolia (Li et al. 1984), the Erlian basin of northern China (E, Fig. 1; Lin et al. 2001), the East Gobi basin in SE Mongolia (G, Fig. 1; Graham et al. 2001; Johnson et al. 2001; Johnson 2004), and the Late Jurassic– Early Cretaceous halfgrabens that exist beneath the extensive Cenozoic post-rift fill of the Songliao basin (S, Fig. 1; Xue & Galloway 1993; Dou 1997). As a whole, this late Mesozoic extensional province is comparable in size and morphology with the modern Basin and Range extensional province of the western USA (Ren & Li 1998). The timing of extension in NE China and Mongolia remains controversial. Extensional basin development appears to have begun as early as Late Jurassic (c. 155 Ma) in the Gobi basin of southern Mongolia (Graham et al. 2001; Johnson et al. 2001; Johnson 2004), in the Erlian basin of Inner Mongolia (Lin et al. 2001), and in Songliao basin (Xue & Galloway 1993), although the rift-related setting of Late Jurassic rocks in this basin remains in question (Wang et al. 2002). Rifting in the Erlain and Gobi basins occurred in three phases, the youngest of which is roughly time-correlative with Early Cretaceous rifting further south (Graham et al. 2001; Lin et al. 2001). With the possible exception of Songliao basin, rifting did not occur throughout most of NE China until Early Cretaceous time, and at least the southernmost part of this region (i.e. the Yanshan belt) experienced contratile tectonism throughout the Jurassic (Davis et al. 1998, 2001). Rifting thus progressed as a sweep of extension that began in northern regions and migrated southward over time. In the Liaoxi region, Cretaceous volcanic strata of variable thickness that range in age from 135 to 124 Ma (Fig. 2) are overlain by fluvial, alluvial, and lacustrine strata that constitute the synextensional sedimentary basin-fill of all the basins described in this paper. It is unclear whether the volcanic rocks that underlie these strata are synextensional in nature, or if they were simply preferentially preserved within downthrown blocks in the half-graben basins. Contact relations in seismic profiles and in outcrop suggest the latter. Volcanic strata are variably preserved only

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along the ramp margins of these half-graben, and do not appear to thicken appreciably towards basinbounding normal faults. The sedimentary fill of Fuxin basin, for example, overlies a much more widepread sequence of Lower Cretaceous volcanic and lacustrine strata (the Yixian Formation) that are well known for housing the feathered dinosaurs of Liaoning Province and dated at 124 Ma (Swisher et al. 1999). These rocks constitute a relatively flat-lying blanket of sediments that cover the Jurasssic–Cretaceous foreland basin sequence over a broad area centred on Yixian (Y, Fig. 3), and their distribution does not appear to have been controlled by the occurrence of extensional structures. Subsidence during this time may have been generated by large-scale, pre-rift magma withdrawal from highly evolved upper crustal reservoirs, as evidenced by the widespread occurrence of rhyolitic volcanism. At least one caldera of probable Early Cretaceous age exists in the Liaoxi region (Fig. 2). This ring-shaped structure bounds a body of hypabyssal rhyolitic– dacitic volcanics c. 25 km in diameter. The best estimate for the age of extension in the Liaoxi region is provided by the ages of sedimentary rocks that are clearly related to subsidence controlled by basin-bounding structures. These units are all younger than 124 Ma, and regionally are correlated with the Lower Cretaceous Jiufotang Formation. Near Dachengzi, biotite from a tuff interbedded with SE-tilted basal lacustrine strata of Jiufotang basin yields an 40Ar/39Ar weighted mean plateau age of 115.1 + 0.7 Ma (Fig. 4b). North of Chaoyang, tuff units interbedded with fossil-bearing Jiufotang lacustrine shales yield 40 Ar/39Ar K-feldspar plateau ages of 120.3 + 0.7 Ma (He et al. 2004). Although these ages do not necessarily constrain the timing of the onset of extension, they do indicate that extensional basins continued to form and fill at least until 115 Ma. Localized high-strain extension resulting in the formation of metamorphic core complexes throughout the entire North China region (Fig. 1) is also Early Cretaceous in age. The lower plate of the Yunmeng Shan metamorphic core complex in the southwestern Hebei segment of the Yanshan belt experienced rapid cooling at c. 118 –115 Ma, probably related to rapid tectonic unroofing along the overlying detachment fault (Davis et al. 1996). On the basis of 40Ar/39Ar cooling ages in footwall mylonites, the Yagan– Onch Hayrhan metamorphic core complex of southern Mongolia and North– Central China experienced extension c. 129– 126 Ma (Webb et al. 1999), and the Hohhot metamorphic core complex was active between 125 and 121 Ma (Davis et al. 2002). The Waziyu metamorphic core complex, which is kinematically

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linked to the formation of the Fuxin supradetachment basin, was probably active between 127 and 116 Ma (Zhang et al. 2003; Darby et al. 2004).

Drivers of extension in NE China and Mongolia The broad spatial and temporal overlap of extension in NE China and SE Mongolia suggests a common tectonic driver. Previous workers have suggested Asia –Pacific plate interactions (Li et al. 1984; Watson et al. 1987; Dou 1997), post-orogenic collapse (Traynor & Sladen 1995; Johnson et al. 2001; Meng 2003), or a combination of these factors (Graham et al. 2001) as possible mechanisms for Late Jurassic–Early Cretaceous extension in NE China and SE Mongolia. Lithospheric foundering has been cited as a possible mechanism for the widespread adakitic magmatism that accompanied Late Jurassic–Early Cretaceous tectonism in the North China region (Gao et al. 2002, 2004), as has palaeo-Pacific subduction (Davis et al. 2001; Davis 2003). Given the association of metamorphic core complexes superimposed upon older sites of crustal thickening and the localized reactivation of thrust structures as normal faults, gravitational collapse seems an attractive mechanism for extension in the region. Overall, the surface and subsurface geology of the Liaoxi region resembles an orogen that has had its roots pulled out from under it by lower crustal flow, resulting in core complex development to the east of Fuxin basin and collapse of orogenic highlands to the NW. An alternative (or co-operative) driver for late Mesozoic rifting in NE China and Mongolia could be a reduction in the rate of Pacific– Eurasia plate convergence, cited as a possible cause of Cenozoic extension in the region (Northrup et al. 1995; Ren et al. 2002). Reconstructions of convergence rates between Eurasia and Pacific basin oceanic plates are problematic because of a lack of evidence as to which oceanic plate was converging with Eurasia at that time (Engebretson et al. 1985). Reconstructions by Engebretson et al. (1985), however, do show an abrupt decrease in convergence rates (from c. 300 to 200 mm a21 with no appreciable change in direction) between the ancestral Pacific plate (Izanagi plate) and Eurasia at 125 Ma, coincident with the onset of extension in NE China. A similar drop in convergence rates (from 102 to 17 mm a21) occurred between Eurasia and the Farallon plate at c. 119 Ma. One or both of these oceanic plates was probably subducting beneath Eastern Asia during Cretaceous time. Hence, a reduction in convergence rates may have

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facilitated extension parallel to the continental margin in NE China during Cretaceous time. The spatial coincidence of Early Cretaceous NW–SE extension with NE-trending sinistral strike-slip is characteristic of NE China, and has been noted by many workers (Davis et al. 2001; Han et al. 2001; Zhang et al. 2003). In the Liaoxi region, sinistral transpression appears to be partially responsible for the inversion of Cretaceous extensional basins, and may reflect reorganization of the palaeo-Pacific plate margin in response to shifting plate motions.

Implications for lithospheric delamination in North China Kimberlite-hosted xenoliths of Ordovician age from the eastern part of the North China block indicate the existence of thick (c. 180 –200 km) lithosphere beneath North China during this time (Fan et al. 2000). Basalt-hosted xenoliths of Cenozoic age indicate a thin (c. 80 km) lithosphere underlain by depleted asthenospheric mantle (Zheng et al. 2005), indicating that the lithosphere in Eastern China was significantly thinned prior to Cenozoic time. Gao et al. (2004) suggested that lithospheric foundering occurred during the late Mesozoic, as indicated by the occurrence of 159 Ma adakitic rocks that have geochemical signatures consistent with derivation by partial melting of Archean lower crustal eclogite. Davis (2003) argued that delamination of the North China lithosphere, if it occurred, probably took place during Late Cretaceous–Cenozoic time, and attributed Jurassic – Early Cretaceous magmatism in the Yanshan region to palaeo-Pacific arc magmatism interacting with thick continental crust. Still another possibility is that magmatism was initiated by post-collisional slab breakoff along the south-dipping subduction zone that existed prior to collision of the Mongolian arc terranes with North China, forming the Suolon suture, or along the north-dipping collision zone between North and South China. It is likely that all of these factors contributed in some fashion to the vast magmatic province that developed in North China during Mesozoic time. Several geologic aspects of the North China block are consistent with late Mesozoic crustal weakening in eastern North China. The Yanshan belt and surrounding regions underwent virtually continuous tectonism and magmatism commencing in Triassic– Early Jurassic time and continuing into the Cenozoic. The Ordos plateau, which occupies a large portion of the western North China block, was stable during the same time period, despite considerable tectonic activity along its margins (Darby & Ritts 2002). The two regions have a

similar stratigraphic history throughout the bulk of the Paleozoic. The entire North China block is characterized by a regional Silurian– Devonian disconformity (Yin & Nie 1996), and Silurian strata are present only near the northern periphery of North China (Cope et al. 2005). The occurrence of this disconformity supports the notion that both Ordos and eastern North China were elevated and stable at least into the Carboniferous. From Triassic time onward, however, the tectonic and stratigraphic histories of these two regions diverge dramatically. Mesozoic strata throughout central Ordos are flat-lying and undisturbed, whereas strata of the same age in eastern North China are highly deformed and sedimentation occurred only within restricted, structurally controlled basins (Cope 2003). This suggests that a profound discontinuity formed between the two regions during Mesozoic time, even though they occupy the same Archaean craton and have similar tectonic histories. The voluminous magmatism that accompanied this event affected an enormous region in NE China and SE Mongolia during the Mesozoic, and is difficult to explain by palaeo-Pacific subduction alone (Meng 2003). Lithospheric thinning would be expected to produce a thermal anomaly that, once decayed, would result in widespread thermal subsidence (McKenzie 1978), particularly if lithospheric mantle is thinned over a broader area than the crust (White & McKenzie 1988). A curious facet of Cretaceous extension in NE China and Mongolia is that all of the basins formed during this event (with the exception of Songliao basin) appear to lack a subsequent post-rift fill (Li et al. 1984; Graham et al. 2001; Lin et al. 2001; Ren et al. 2002; Meng et al. 2003), which is difficult to reconcile with the amount of extension and magmatism that occurred throughout the region (Meng 2003), and even more difficult to reconcile with lithospere loss. Like the basins of the Liaoxi region, basins in southern Mongolia experienced a post-rift episode of inversion that was possibly related to left-lateral transpression (Graham et al. 2001), and may have terminated further subsidence. Synrift sequences in the East Gobi basin were later overlain by a thin, Late Cretaceous cover (Graham et al. 2001) that may represent a late effect of post-rift thermal subsidence. The marginal Cenozoic sedimentary basins of Eastern China, however, do contain significant post-rift fill, despite widespread transpressional inversion. Songliao, North China and Bohai Bay basins (Fig. 1) all consist of a faulted synrift sequence that is overlain by relatively unfaulted ‘sag phase’ strata (Watson et al. 1987; Hsiao et al. 2004). All of these basins lie along the trend of the Tan-Lu fault (Fig. 1), and this fault has

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been shown to be operational during the Cenozoic, resulting in transpressional and transtensional deformation that did not affect the subsequent development of post-rift subsidence within these basins (Hsiao et al. 2004). Inversion therefore seems an insufficient mechanism to account for the lack of such subsidence, particularly given the large thermal anomalies that must have existed in eastern North China during the Mesozoic. The lack of post-rift subsidence in Cretaceous basins throughout China and Mongolia could be attributable to the mitigating effect of middle– lower crustal flow that regionally smoothed thermal anomalies across the rift province and suppressed post-rift thermal subsidence, as suggested by Meng (2003). This theory is consistent with the widespread occurrence of metamorphic core complexes throughout the region, which indicate mid- to lower crustal mobility during rifting, and supports the notion that gravitational collapse above thermally weakened and overthickened crust was the primary driver for mid-Cretaceous extension. If lithosphere removal indeed occurred during Mesozoic time, it was probably only a local phenomenon; wholesale thinning of the lithosphere is more consistent with the broad zones of thermal subsidence that formed in eastern North China during the Cenozoic.

Conclusions Extensional basins developed within the Liaoxi region share many characteristics with coeval basins throughout NE China and southern Mongolia, in that they lack post-rift thermal subsidence, are generally bounded by low-angle ductile and brittle normal faults that root at mid-crustal levels, and accommodate NW–SE-directed extension. The basins in Liaoxi commonly are bounded by normal faults that sole into older thrusts, a geometry that may be shared by other Cretaceous basins in the extensional domain of Eastern Asia. The similarity of these basins, in terms of timing, extension direction and structural style, suggests a regional tectonic driver. Most of the region that extended during Early Cretaceous time is was underlain by thickened crust that was subject to extensive, and largely adakitic, magmatism prior to (and possibly during) extension. In the Yanshan region, this magmatism commenced as early as 219 Ma, broadly coeval with plate collision along both the southern and northern margins of North China. It is unclear whether magmatism is related to lithospheric removal, slab breakoff along suture zones to the north or south, palaeo-Pacific subduction, or a combination of all of these factors.

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Although it is clear that lithospheric delamination did occur at some point during the history of the North China block, there appears to be no temporal link between postulated Mesozoic delamination (as old as 160 Ma, Gao et al. 2004) and the widespread extension that characterizes NE China and SE Mongolia during mid-Cretaceous time (post-130 Ma). In addition, the lack of post-rift subsidence throughout the Cretaceous rift province described in this paper, and the presence of such subsidence in Cenozoic basins, points to a later period of lithospheric thinning. Hence, although thermal weakening of the lower crust during Mesozoic magmatism may have facilitated Cretaceous extension by initiating lower crustal flow outward from thickened regions, there is little upper crustal evidence for wholesale lithospheric thinning during Cretaceous time. Extension was probably driven by a combination of plate boundary forces and internally driven gravitational collapse, facilitated by thermal weakening of the crust as a result of c. 219–130 Ma magmatism. The authors are deeply indebted to many people for making this work possible. This research was enhanced as the result of lively debates with G. A. Davis, who has consistently regarded our research in NE China with a healthy dose of scepticism. We are greatly indebted to Liang Hongde for showing the lead author key outcrops and field relations in Liaoxi. This work benefited greatly from discussions with J. Mills and M. Scott Wilkerson at DePauw. Li-Yuan Hsiao was a helpful companion and friend during fieldwork and data mining in China. A. Egger, Xiao Weifeng, Sun Weihua, and A. Gehlhausen all contributed their insight and field skills to various parts of this endeavour. E. Chang and Zhang Changhou provided valuable ground support in China. Seismic data were provided by an anonymous donor. This research was funded by the Stanford–China Industrial Affiliates Program, the DePauw University Professional Development Fund, and a DePauw Freeman Faculty Scholarship.

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Phanerozoic evolution of the Inner Mongolia– Daxinganling orogenic belt in North China: constraints from geochronology of ophiolites and associated formations L. MIAO1, F. ZHANG1, W.-M. FAN1 & D. LIU2 1

Institute of Geology and Geophysics, Chinese Academy of Sciences, Beijing 100029, China (e-mail: [email protected]) 2

Beijing SHRIMP Center, Institute of Geology, Chinese Academy of Geosciences, Beijing 100037, China Abstract: The Inner Mongolia– Daxinganling Orogenic Belt (IMDOB), located between the North China and South Mongolia Blocks, consists of several ENE –WSW- to NE– SW-trending zones including dismembered ophiolite blocks, metamorphic rocks and granitoids. Although numerous studies have been carried out on this belt, its tectonic evolution has been a subject of controversy, chiefly because of the lack of reliable geochronological data. Based on a synthesis of newly published geochronological data and our unpublished data for the IMDOB, we define two oceanic basins: Ondor Sum and Hegenshan. The former, probably the main one, was initiated during the Ordovician (.467 Ma) period, whereas the latter, representing a back-arc basin, opened on a pre-Permian basement at, or earlier than, Early Permian times (c. 295 Ma). These two oceanic basins were separated by a magmatic arc (Sunid–Baolidao), and were probably closed simultaneously when the final orogenesis of the IMDOB occurred during the Triassic period (240– 220 Ma). Importantly, the Triassic timing of the final orogenesis of the IMDOB due north of the North China Craton is essentially coeval with that of the Qinling– Dabie–Su– Lu orogenic belt on the southern margin of the North China Craton. It is inferred that this two-sided subduction– collision scenario in the Triassic may have contributed to the Mesozoic lithospheric thinning event of the North China Craton, although the details are unclear.

The Inner Mongolia –Daxinganling Orogenic Belt (IMDOB), the Eastern Chinese extension of the Central Asian Orogenic Belt (Zonenshain 1973) (the equivalent to the Manchurides and the southern portion of the Alaids of Sengo¨r & Natal’in (1996) or to the Neimonides of Hsu¨ et al. (1991)), resulted from closure of the Palaeo-Asian ocean(s) between the Sino-Korean Craton and the South Mongolia Block (Fig. 1a; Mossakovsky et al. 1994; Jahn et al. 2000, 2004; Badarch et al. 2002). Although numerous studies have been performed on this belt in efforts to delineate the architecture and to reconstruct the evolutionary history of the orogenic belt (Wang & Liu 1986; Li 1987; Zonenshain et al. 1990; Tang & Yan 1993; Sengo¨r & Natal’in 1996; Chen et al. 2000; Wu et al. 2002; Xiao et al. 2003, and references therein), no consensus on its tectonic evolution, especially the timing and location of the final suturing, have been reached. Wang & Liu (1986) proposed a model of accreted terranes. In their model, the terrane accretion was presumed to take place on both the northern margin of the SinoKorean Block and the southern margin of the Siberian (South Mongolia) Block. They argued that the final collision of the two blocks in central Inner

Mongolia occurred along a line from Solon Obo to Linxi at the end of the Permian. On the basis of a comparison with Alpine tectonic facies, Hsu¨ et al. (1991) suggested a tectonic model for the IMDOB (called the ‘Neimonides’). According to this model, the ophiolites are pieces of large me´lange zones comprising disrupted sedimentary– volcanic and metamorphic rocks, as well as fragments of ophiolites. Hsu¨ et al. (1991) identified two ophiolitic me´langes, Ondor Sum and Hegenshan, and interpreted them to represent relicts of two separate Palaeozoic oceans or basins, with the Ondor Sum ocean being closed during the early Palaeozoic and the Hegenshan during the late Palaeozoic. In the model, a micro-continental block (Xilinhot) was postulated to separate the two oceans and to be welded to the northern margin of the Sino-Korean Block when the Ondor Sum ocean closed. Robinson et al. (1999) modified this model, and they suggested an island arc rather than a micro-continent to separate the two oceans. Tang & Yan (1993) put forward a tectonic model for the IMDOB (they named it the ‘Inner Mongolian suture zone’) based on regional metamorphism. They recognized four stages of metamorphism during the Palaeozoic, and related them

From: ZHAI , M.-G., WINDLEY , B. F., KUSKY , T. M. & MENG , Q. R. (eds) Mesozoic Sub-Continental Lithospheric Thinning Under Eastern Asia. Geological Society, London, Special Publications, 280, 223–237. DOI: 10.1144/SP280.11 0305-8719/07/$15 # The Geological Society of London 2007.

Fig. 1. Geological and tectonic sketch map of the Inner Mongolia–Daxinganling Orogenic belt (IMDOB) (modified from Wang & Liu 1986; GMRBIM 1991). The inset shows the position of the IMDOB in East Asia (modified from Sengo¨r & Natal’in 1996). Numbers represent ophiolites and arc or metamorphic complexes (see Table 1). Dates for these major bodies are also indicated, and abbreviations in parentheses following the age indicate the method used (Rb, Rb –Sr isochron; Ar, Ar– Ar method; SH, SHRIMP zircon U –Pb method). EH, Eren Hot; HG, Hegenshan.

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to subduction, collisional and/or post-orogenic events. Their model postulated that two oceanic basins sequentially opened and closed (Ondor Sum and Hegenshan) in the IMDOB. The Ondor Sum ocean existed between the North China Block (NCB) and the South Mongolia Block (SMB) from the Late Proterozoic to Cambrian, the subduction of which during or before the Ordovician generated an island arc (Bainaimiao) and a continental magmatic arc (Dong Ujumqi) on the northern margin and southern margin of the NCB and SMB, respectively. The Hegenshan ocean opened during the Early Silurian, contemporaneous with, or slightly later than, the closure of the Ondor Sum ocean. In their model, the newly opened Hegenshan ocean began its subduction during the earliest Devonian and closed during the Late Devonian when the two continental blocks collided. They further suggested that the latest metamorphism, which occurred during the Late Palaeozoic, was related to post-collisional rifting in the region. Sengo¨r & Natal’in (1996) divided the mountain belts between the Sino-Korean and Siberian Cratons into the Altaids and Manchurides, which represent collages of subduction–accretion complexes accreted onto the southern and the northern margins of the Siberian and the Sino-Korean Cratons, respectively. They suggested that such subduction–accretion complexes were built by suturing small pieces of lithosphere across subduction zones to the opposing continental margins. According to this model, the Altaids and Manchurides were welded together along the Solonker suture when the ocean (Solonker) between them was closed in the Late Permian. Nozaka & Liu (2002) suggested that the Hegenshan ophiolite represents the location of the final suturing of the IMDOB, and inferred, based on two amphibole K– Ar dates, a Cretaceous suturing time. Xiao et al. (2003) proposed a model similar to that of Sengo¨r & Natal’in (1996), but provided more details. They suggested that the CAOB underwent three evolutionary stages in its termination history: an Early to Mid-Palaeozoic stage of accretion of subduction–accretion complexes to the opposing margins, a Permian stage of Andean-type margins developed on the consolidated accretionary complexes, and an end-Permian stage of Himalayantype collision caused by two-way subduction beneath the Andean-type margins. The variety of models described can partially be attributed to the complexity of the IMDOB itself and the scarcity of outcrops as a result of the low relief and abundant steppe vegetation, but the paucity of reliable geochronological evidence is also a factor. This paper synthesizes recently published and unpublished reliable geochronological

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data (e.g. zircon sensitive high-resolution ion microprobe (SHRIMP) and Ar –Ar dates) for the IMDOB, and discusses its tectonic evolution based on these data. Additionally, the possible influence of the orogenesis on the North China Craton (NCC) is discussed.

Regional setting Within central Inner Mongolia, the IMDOB is composed of a wide orogenic belt, containing numerous ophiolitic blocks with variable ages (Wang & Liu 1986), which are broadly subdivisible into four NE –SW- to ENE– NSW-trending zones, form south to north, namely Ondor Sum –Xar Moron, Solon Obo –Linxi, Jiaoqier –Xilinhot, and Eren Hot –Hegenshan (Fig. 1b). The Ondon Sum– Xar Moron zone extends from north of Baiyun Obo in the west, eastward via Ondor Sum, to the Xar Moron River in the east. This zone comprises the Ondor Sum ophiolite in the central and several ultramafic –mafic blocks in the east. The Ondor Sum ophiolite and its surrounding rocks, called the Ondor Sum or Wenduermiao Group, have been metamorphosed to greenschist and blueschist facies (Tang et al. 1983; Wang & Liu 1986). The ultramafic –mafic blocks scattered along the northern bank of the Xar Moron River in the east of the zone are composed of metamorphic peridotite, gabbro, and/or mafic lava, and are in fault contact with the surrounding low-grade metamorphic rocks of early and middle Silurian age (Wang & Liu 1986). Just to the south of the Ondor Sum ophiolite is the Early Palaeozoic Bainaimiao arc containing calc-alkaline tholeiitic basalts to minor felsic lavas, alkaline basalts, and agglomerates, tuffs, granodiorites and granites. The arcrelated granodiorites have high initial strontium isotope ratios of 0.7146 (Shao & Zhang 1998) and relatively low 1Nd(t) values of 2.4 + 1.7 (Nie & Bjørlykke 1999), probably suggesting formation in an Andean-type continental margin setting. The Solon Obo –Linxi ophiolite belt, extending from the China–Mongolian border in the west, easterly through Sunidyouqi, to Linxi in the east, contains ophiolitic complexes in its western, central, and eastern portions. In the western end, the ophiolite suite (Solon Obo) consists of harzburgite, dunite, gabbroic cumulate, pillow lavas, and chert, and they are covered by sediments of Lower Permian age. Carboniferous coral and brachiopod fossils are found in the surrounding sedimentary strata (Wang & Liu 1986). In the central section of the belt (c. 15 km NE of Sunidyouqi), ophiolites, as well as cherts, marble and island-arc volcanic rocks, occur as exotic blocks within an argillite matrix, and these have

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been considered to represent a tectonic me´lange (Tang & Yan 1993). This me´lange is in fault contact with a turbidite unit to the north. Only a few fossils of Tabulata are found in a limestone lens (Tang 1990). In the western section of the belt, near Linxi, ophiolitic lenses occur in Permian clastic sediments as well as Carboniferous shales and siltstones of the Linxi Formation, and comprise pyroxenite, layered gabbro, sheeted mafic dykes, basalt and chert that contains Early Palaeozoic conodonts such as Panderodus sp. (Wang & Liu 1986). The Jiaoqier– Xilinhot zone extends from Jiaoqier (c. 30 km south of Sunidzuoqi) in the west, through Xilinhot, to south of Xi Ujumqinqi, and is characterized by ophiolites, metamorphic complexes and blueschists, as well as granites. To the south of Sunidzuoqi, the Jiaoqier ophiolitic complex is composed of serpentinized peridotite of Alpine type (Wang & Liu 1986), diabase, pillow lavas, and plagiogranite, and is in fault contact with the surrounding rocks of Carboniferous– Early Permian age. The metamorphic complex, known as the Xilin Gol Complex, is mainly composed of biotite and muscovite gneisses and amphibolites, with intercalated marble and metachert. This metamorphic complex has long been inferred to represent a Precambrian micro-continental block by nearly all researchers, except for Tang (1990), who considered it a metamorphosed Early Palaeozoic flysch formation. The blueschists occur west of the Jiaoqier ophiolite. Two types of granites, arc- and collision-related ones, have been identified in the zone (Chen et al. 2000). In addition, this zone has been undergone intensive deformation, which is characterized by a north-directed nappe structure, from which muscovite yielded K –Ar ages of c. 245 Ma (Zhang et al. 1995). Ultramafic –mafic blocks also occur at c. 15 km south of Xi Ujumqinqi and are associated with Carboniferous –Permian volcano-sedimentary sequences. The Eren Hot– Hegenshan zone, from Eren Hot near the China –Mongolia border in the SW to Hegenshan in the NE, contains numerous ophiolite slices. These occur in a flysch formation of Carboniferous –Permian age (Wang & Liu 1986; Liang 1991). In the Hegenshan region, the ophiolite sequence has experienced intensive deformation and metamorphism, but unmetamorphosed basalt, limestone and chert also exist. Permian carbonbearing shales and siltstones (the Linxi Formation) occur in the eastern extension of this belt. To the north of the Eren Hot–Hegenshan ophiolite zone is the Uliastai (or Dong Wjumqin) belt, characterized by active continental marginal formations composed of large volumes of basalts, andesites and pyroclastic deposits of Carboniferous– Permian age. These volcano-sedimentary sequences overlie Cambrian –Ordovician– Silurian– Devonian

shallow marine and continental fine clastic rocks. This belt was intruded by A-type granitic plutons of Late Carboniferous–Early Permian age. These A-type granites were considered to be postcollisional by Hong et al. (1995), whereas Wu et al. (2002) pointed out that their emplacement was coeval with subduction of the ocean. Immediately south of the Uliastai belt, a Carboniferous – Early Permian accretionary wedge comprising strongly deformed interbedded hemipelagic shales, turbidites and conglomerates with pebbles of chert has been recognized (Hsu¨ & Chen 1999; Xiao et al. 2003). These ophiolite belts and the Palaeozoic formations were covered or intruded by Jurassic– Cretaceous (Yanshanian stage) acid to intermediate igneous rocks.

Geochronological data Geochronological data for the IMDOB in central Inner Mongolia, China, are summarized in Table 1, and are briefly described below.

Ondor Sum – Xar Moron belt The Ondor Sum– Xar Moron belt contains several ophiolite blocks, such as the Ondor Sum, Kedanshan, Banlashan, and Xinshuwa. Of these the Ondor Sum ophiolite is the largest, and has been well studied. The Ondor Sum ophiolite, including two disrupted slices of Ondor Sum and Tulinkai, was considered to be a component of a Palaeozoic me´lange (Hsu¨ et al. 1991; Xiao et al. 2003). The geology of the me´lange as well as the Ondor Sum ophiolite has been described by Xiao et al. (2003). The matrix of the me´lange, called the Ondor Sum Group, has been deformed and metamorphosed to form greenschist and blueschist. Metabasic volcanic rocks (greenschists) have contradictory Sm – Nd and Rb–Sr isochron ages of 961 + 66 and 624 + 110 Ma, respectively (Zhang & Wu 1998), and glaucophane from the blueschists has been dated by Ar –Ar isochron method at 426 + 15 Ma (Tang & Yan 1993). Cumulate gabbroic rocks from the Tulinkai slice have a zircon SHRIMP age of 457 + 4 Ma (Fig. 2a); a diorite intrusion with an adakitic geochemical signature in the area has a zircon SHRIMP age of 467 + 6 Ma (Liu et al. 2003). These dates suggest an early Palaeozoic formation time for the Ondor Sum (Tulinkai) ophiolite. However, a pillow lava sample without deformation from the Ondor Sum slice gave a zircon SHRIMP age of c. 260 Ma, although older ages of c. 440 Ma were also detected (Fig. 2b). This indicates that the pillow lava probably formed later than the cumulate gabbros.

Ondor Sum Group

Hegenshan ophiolite (12)

Granodiorite Granodiorite Greenschist Granodiorite Pl-crystal tuff Metabasalt Metabasalt Metabasalt Blueschist Pillow lava Cumulate gabbro Diorite Ultramafic rock Cumulate gabbro Cumulate gabbro Biotite gneiss Cumulate gabbro Diorite Cumulate gabbro Plagiogranite Blueschist Biotite gneiss Cumulate gabbro Tectonic schist Quartz diorite Quartz diorite Granite Potassic granite Mafic – ultramafic Metamafic dyke Ultramafic rock Cumulate gabbro Mafic dyke Mafic lava Granodiorite dyke Biotite gneiss Bi-mica schist

Measured rock WR Zr WR Zr Zr WR WR WR Gl Zr Zr Zr WR Zr Zr Zr Zr Zr WR Zr Gl Zr Zr Mus Zr Zr WR Zr WR Hb WR Zr Zr WR Zr Zr WR

Measured object Sm – Nd isochron Conventional U– Pb Rb – Sr SHRIMP SHRIMP Sm – Nd Rb – Sr isochron Rb – Sr isochron Ar– Ar SHRIMP SHRIMP SHRIMP K– Ar SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP Rb – Sr isochron SHRIMP Ar– Ar SHRIMP SHRIMP K– Ar SHRIMP SHRIMP Rb – Sr isochron SHRIMP Sm – Nd isochron Ar– Ar K– Ar SHRIMP SHRIMP Ar– Ar SHRIMP SHRIMP Rb – Sr isochron

Method

Data source Nie & Bjørlykke (1999) Tang & Yan (1993) Tang & Yan (1993) This paper This paper Zhang & Wu (1998) Zhang & Wu (1998) Tang (1990) Tang & Yan (1993) This paper This paper Liu et al. (2003) Liang (1991) This paper This paper G. H. Shi (unpubli. data) This paper This paper Wang & Liu (1986) This paper Xu et al. (2001) Shi et al. (2003) G. H. Shi (unpubli. data) Zhang et al. (1995) Chen et al. (2000) Chen et al. (2000) Chen et al. (2000) Shi et al. (2004) Bao et al. (1994) Robinson et al. (1999) Liang (1991) This paper This paper This paper This paper Miao et al. (2004) Wu et al. (2004)

Age (Ma) 429 + 100 466 428 + 17 430 + 7 440 + 7 961 + 66 624 + 110 509 + 40 426 + 15 c. 260 457 + 4 467 + 6 c. 570 73 + 2 256 + 3 270 279 + 10 438 + 4 262 343 + 7 383 + 13 437 + 3 256 245 309 + 8 490 + 8 228 + 21 220 – 200 403 + 27 244 + 2 346 – 600 295 + 15 298 + 9 293 + 1 244 + 4 216 + 2 240

OX, Ondor Sum – Xara Moron belt; SL, Solon Obo –Linxi belt; JX, Jiaoqier –Xilin Hot belt; EH, Eren Hot –Hegenshan belt; WR, whole rock; Zr, zircon; Gl, glaucophane; Mus, muscovite; and Hb, hornblende. *Numbers correspond to those in Figure 1.

Xinkaling Complex in NE China Hulan Group in NE China

EH

Ondor Sum ophiolite (2-1) Tulinkai (2-2) Tulinkai Kedanshan complex (3) Kedanshan complex (3) Banlashan ophiolite (4) Shuangjing complex (5) SL Solun Obo ophiolite (6) Solun Obo Balengshan ophiolite (7) JX Jiaoqier ophiolite (8) Jiaoqier blueschist Xilin Gol complex (9) Xilin Gol ophiolite (10) Jiaoqier nappe Sunid arc (11)

OX

Bainaimiao arc (1*)

Belt Formation

Table 1. Summary of geochronological data for the IMDOB

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Fig. 2. Concordia plots showing SHRIMP zircon U–Pb analytical data for rocks from the Ondor Sum–Xar Moron belt in central Inner Mongolia. (a) A cumulative gabbro (sample 2002MW-01) from the Tulinkai ophiolitic block (No. 2-2 in Fig. 1 and Table 1); (b) a pillow basalt sample (sample 2002MW-03) from the Ondor Sum ophiolite (No. 2-1 in Fig. 1); (c) a cumulative gabbro sample (02NMG-14) from the so-called Kedanshan ophiolite (No. 3 in Fig. 1), which yielded a very young age of c. 73 Ma; (d) a cumulate gabbro sample (2002ML06-4) from the Banlashan ophiolite (No. 4 in Fig. 1). Light grey error ellipses (1s) represent analytical spots included in age calculation of the samples; dark grey error ellipses are those not included in this study.

In the eastern portion of the belt, ultramafic – mafic blocks are scattered along the northern bank of the Xar Moron River (Fig. 1). The so-called Kedanshan ophiolite is nearly entirely covered by Quaternary sediments, and ultramafic –mafic rocks can be observed only in man-made trenches. These rocks have long been considered to be early Palaeozoic in age, chiefly based on a single whole-rock K– Ar age of c. 570 Ma (pers. comm., cited by Liang 1991). However, our new zircon dating on the cumulate gabbro gave an extremely young age of 73 + 2 Ma (Fig. 2c), although old zircons with ages between 107 and 2536 Ma were also detected. Because the young age of 73 + 2 Ma was obtained from zircons with typical magmatic zonation (Fig. 3a), we interpreted it as the emplacement age

of the gabbro. The zircons yielding the older ages are mostly rounded grains without zonation, and are interpreted to be xenocrystal or inheritance in origin. Consequently, we preliminarily believe that the Kedanshan ultramafic–mafic body is probably not ophiolitic but intrusive in origin. This postulation is also supported by our field observations, and petrological and geochemical data: for example, the existence mafic dykes probably derived from the body intruding the Permian volcano-sedimentary sequence (Fig. 3b); the cumulate rocks being characterized by containing noritic gabbros, which are rare in typical ophiolites; and the Nd isotope composition being characterized by 1Nd(t) values between 20.6 and þ1.2 (L. C. Miao, unpubl. data). Cumulate gabbro from another ultramafic–mafic body

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Fig. 3. (a) Representative cathodoluminescence (CL) images of zircons from a cumulate gabbro of the Kedanshan ultramafic–mafic complex (SHRIMP analytical spots and corresponding ages are also indicated); (b) a photograph showing a mafic dyke intruding a Permian clastic sequence near the Kedanshan ultramafic– mafic complex.

(Banlashan) along the Xar Moron River has a zircon SHRIMP U–Pb age of 256 + 3 Ma (Fig. 2d). Apart from the country rocks of the Ondor Sum ophiolite, metamorphic rocks locally occur along the northern bank of the Xar Moron River, where the rocks were collectively named the ‘Shuangjing Complex’ (Geology and Mineral Resource Bureau of Inner Mongolia (GMRBIM) 1991). The complex mainly consists of biotite–muscovite gneisses, with minor amount of marbles and schists. Although no age and fossil data are available, the Shuangjing Complex has been thought to be Proterozoic or Ordovician in age (BGMRIM 1991; Shao 1991). Zircons from a mica gneiss of the complex have an SHRIMP age of c. 270 Ma (G. H. Shi, unpubl. data).

Bainaimiao arc Located due north of the NCC, the Bainaimiao arc contains calc-alkaline tholeiitic basalts to minor

felsic lavas, alkaline basalts and agglomerates, volcanic breccias, tuffs, granodiorites and granites. These rocks have been metamorphosed at a greenschist-facies condition (Tang & Yan 1993). Granodiorites from this arc have conventional zircon U – Pb and whole-rock Sm –Nd isochron ages of 466 Ma and 429 + 100 Ma, respectively (Tang & Yan 1993; Nie & Bjørlykke 1999), and a greenschist has a whole-rock Rb–Sr isochron age of 428 + 17 Ma. Our new zircon SHRIMP dates are broadly consistent with these results. A greenschist from the arc formation gave an age of 440 + 7 Ma, along with several zircons giving relatively young ages (Fig. 4b), and a granodiorite yielded a SHRIMP age of 430 + 7 Ma (Fig. 4a), although older ages of c. 520 Ma were also obtained. The presence of zircon inheritance of the granodiorite is in agreement with its Sr –Nd isotope signature, which is characterized by a high initial strontium isotope ratio (0.7146; Shao 1991)

Fig. 4. Concordia plots showing SHRIMP zircon U–Pb analytical data for granodiorite (a) and metavolcanic rock (b) from the Bainaimiao arc (No. 1 in Fig. 1), central Inner Mongolia.

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and a relatively low 1Nd(t) value (2.4 + 1.7; Nie & Bjørlykke 1999). These data indicate that the Bainaimiao arc was probably formed on a continental margin during Silurian time.

Solon Obo – Linxi belt The Solon Obo –Linxi belt has been considered by some workers to represent the final suturing zone of the Central Asian Orogenic Belt (e.g. Wang & Liu 1986; Sengo¨r & Natal’in 1996; Xiao et al. 2003), but no detailed dating data are available except for a single whole-rock Rb –Sr isochron age of c. 260 Ma for the Balengshan ophiolite at the eastern end of the belt near Linxi (Wang & Liu 1986). We dated the Solon Obo ophiolite, the largest one in the belt. The Solon Obo ophiolite, which straddles the China –Mongolia border with about a half on each side, is composed of harzburgite, dunite, gabbroic cumulate, mafic pillow lavas and chert, and is overlain by Lower Permian clastic sediments (Wang & Liu 1986). A small diorite intrusion occurs in the southern part of the ophiolite, and has a fault contact with it. A diorite sample from the intrusion has a weighted mean age of 438 + 4 Ma for the main zircon population (Fig. 5a), with one analysis yielding a young age of 307 + 6 Ma. We interpret the age of 438 + 4 Ma age of the main zircon group to be the emplacement age of the intrusion, and the young age probably to be a result of isotopic resetting by late thermal events for the analysed spot. A cumulate gabbro sample of the ophiolite, however, yielded a SHRIMP age of 279 + 10 Ma (Fig. 5b), with some older zircons giving ages of 440 –450 Ma. The older ages of the gabbro are broadly consistent with the emplacement age of the

diorite, and are therefore interpreted to be inherited or xenocrystic in origin. These data suggest that the Solon Obo diorite and ophiolite are not cogenetic. We interpret the former to represent a component of an early Palaeozoic arc formation, and the latter a late Palaeozoic ophiolite. The association with diorite and the zircon inheritance of the Solon Obo ophiolite possibly indicate that the ophiolite formed in a suprasubduction-zone environment.

Jiaoqier – Xilin Hot belt This belt is characterized by containing ophiolite blocks within intensively metamorphosed rocks. The metamorphic rocks, named the Xilin Gol complex, have been thought to be Precambrian in age, chiefly based on unreliable conventional zircon U –Pb ages of c. 1060 Ma (BGMRIM 1991) and Sm– Nd and Rb– Sr ages of c. 1225 and c. 650 Ma (Xu et al. 1996). However, newly published SHRIMP zircon U –Pb data indicate that this metamorphic complex is not older than Silurian in age (Shi et al. 2003). According to Shi et al. (2003), a biotite gneiss from the complex near Xilin Hot city, where the complex was primarily named the ‘Xilin Gol complex’, contains two populations of zircons: one is characterized by typical magmatic oscillatory zoning and yielded a weighted mean age of 437 + 3 Ma (n ¼ 13, MSWD ¼ 1.76), and the other by rounded grains yielding scattered ages between 600 and 3100 Ma. Because the protolith of the gneiss is volcanosedimentary in origin, the younger age of 437 + 3 Ma was interpreted by Shi et al. to be representative of the maximum formation age of the complex. Existence of Precambrian zircons in the gneiss indicates that the previous conventional

Fig. 5. U– Pb concordia showing zircon SHRIMP analytical data for quartz diorite (a) and cumulate gabbro (b) of the Solon Obo ophiolite (No. 6 in Fig. 1), central Inner Mongolia.

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zircon U –Pb ages, as well as Sm–Nd and Rb–Sr ages, are unreliable. Considering an Ar –Ar age of 383 + 13 Ma (Xu et al. 2001) of the high-pressurefacies blueschists in this belt, we interpret the metamorphic complex to be a metamorphosed Silurian– Devonian accretionary wedge. A plagiogranite from the Jiaoqier ophiolite was dated at 343 + 7 Ma (Fig. 6), suggesting a Carboniferous time of formation of the ophiolite. A cumulate gabbro from the Xilin Gol ultramafic– mafic complex (No. 10 in Fig. 1b) gave an age of c. 256 Ma (G. H. Shi, unpubl. data).

depended on the isotopic method used, which has led to ambiguity in the formation and emplacement ages of the ophiolite (Wang & Liu 1986; Bao et al. 1994; Robinson et al. 1999; Nozaka & Liu 2002). Our new SHRIMP and Ar –Ar dating results suggest that the ophiolite formed during the late Carboniferous and early Permian, as indicated by a gabbro from the cumulate sequence and a mafic dyke intruding the mantle harzburgite, which yielded zircon SHRIMP ages of 295 + 15 and 298 + 9 Ma (Fig. 7a and b), respectively, and by the massive basalt, which has a whole-rock Ar – Ar age of 293 + 1 Ma (Fig. 7d). Furthermore, a granodiorite dyke that intrudes the ultramafic rocks has a zircon SHIRMP age of 244 + 4 Ma (Fig. 7c), which is consistent with a previous amphibole Ar–Ar age of 244 + 2 Ma (Robinson et al. 1999). Therefore, we interpret the age of c. 244 Ma to represent the time of peak metamorphism or emplacement time of the ophiolite. It is important to note that Nozaka & Liu (2002) obtained two amphibole K –Ar ages of 110 + 5 and 130 + 6 Ma for the metagabbro of the Hegenshan ophiolite and interpreted them to date the peak metamorphism related to the final suturing of the IMDOB. However, this interpretation, as noted by Zhou et al. (2003), is obviously in contradiction with the geology, which shows the whole IMDOB is covered and intruded by large volumes of intracontinental volcanic rocks and plutons of Late Jurassic to Early Cretaceous age. In our opinion, these Cretaceous ages probably result from isotopic resetting by the coeval magmatic event.

Eren Hot – Hegenshan belt

Sunid – Baolidao arc

The ophiolites near Eren Hot are poorly exposed because of Quaternary sediment cover, whereas those at Hegenshan crop out relatively well. At Hegenshan, more than 20 dismembered ophiolite slices have been identified, which are collectively named the ‘Hegenshan ophiolite’. Drilling and a magnetotelluric study have shown that these slices belong to one giant ophiolite at depth (Hsu¨ et al. 1991). The Hegenshan ophiolite, associated with unmetamorphosed basalt, limestone and chert, is mainly composed of metamorphic harzburgite, dunite, ultramafic and mafic cumulates, and basalt (greenschist), and is intruded by mafic and granodiorite dykes. Additionally, the ophiolite contains significant chromite mineralization. Chromite mineralogy and geochemistry of mafic rocks suggested that the Hegenshan ophiolite formed in a back-arc basin setting, with unmetamorphosed basalt that has an ocean island basalt (OIB)-like origin (Robinson et al. 1995, 1999). Previous dating on the ophiolite has produced inconsistent results ranging from 110 to 600 Ma,

The Sunid magmatic arc, located between the Jiaoqier –Xilin Hot and Eren Hot –Hegenshan ophiolite belts, extends northeasterly from Sunidzouqi in the west, via Xi Ujimqin, to Huolinhe in the east. Near Sunidzouqi, two types of granitoids within the arc have been identified: one is arc-type and the other is collisional-related. The former is composed of variably deformed, metaluminous to weakly peraluminous, hornblende-bearing gabbroic diorite, quartz diorite, tonalite and granodiorite, and two quartz diorites have zircon SHRIMP U –Pb ages of 490 + 8 and 309 + 8 Ma (Chen et al. 2000), respectively, indicating that the arc is a composite one. The latter consists of two-mica adamellite, granodiorite, and leucogranite, which has a Rb–Sr isochron age of 228 + 21 Ma (Cheng et al. 2000) and zircon SHRIMP ages of 220– 200 Ma (Shi et al. 2004), suggestive of a Triassic collision time for the IMDOB. In the eastern segment between Xi Ujimqin and Huolinhe, this arc is characterized by volcanic rocks (Dashizhai Formation) with a geochemical

Fig. 6. Concordia plots showing zircon SHRIMP U– Pb analytical results for plagiogranite from the Jiaoqier ophiolitic block (No. 8 in Fig. 1), central Inner Mongolia.

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Fig. 7. Diagrams showing the isotopic dating results for the Hegenshan ophiolite (No. 12 in Fig. 1). (a) U –Pb concordia for a mafic dyke (sample 02NMG-1); (b) concordia for a leucogabbro (sample 02NMG-5); (c) concordia for a plagiogranite dyke (sample 21DW-25); (d) Ar–Ar age spectrum for the massive basalt (sample 21DW-37) with OIB-like geochemistry.

affinity to typical arc-type (F. Guo, pers. comm.). The Dashizhai Formation, composed mainly of basalt, andesite, dacite, rhyolite, and volcanic breccias, has been assigned to the early Permian and has a whole-rock K –Ar age of c. 250 Ma (Lu¨ et al. 2002), broadly consistent with the emplacement age (c. 310 Ma) of the Sunid arc-type granitoids.

Discussion and conclusions Timing of suturing The timing of the final suturing between the Sino-Korean and the Siberian Blocks, or the Manchurides and Altaids of Sengo¨r & Natal’in (1996), has long been controversial. The proposed suturing time ranges from Devonian to Cretaceous (e.g. Wang & Liu 1986; Tang 1990; Ruzhentsev & Pospelov 1992; Hong et al. 1995; Sengo¨r & Natal’in 1996; Nozaka & Liu 2002; Xiao et al. 2003). New zircon SHRIMP and Ar –Ar dates now provide some important constraints on this issue.

The youngest ophiolites, such as the Banlashan, Solon Obo and Xilin Gol, in the IMDOB are late Permian in age (250 –270 Ma). This indicates that oceanic basin(s) were still developed in that time. Granodiorite dykes intruded the Hegenshan ophiolite at 244 + 4 Ma (earliest Triassic). The geochemistry of these dykes (e.g. very low heavy rare earth elements, positive Sr anomaly, high 1Nd(t) values between þ6 and þ8, and low ISr ratios ranging from 0.70412 to 0.70450; Miao et al. submitted) suggests that they probably formed from partial melting of oceanic materials. Considering the Ar –Ar date of 242 + 2 Ma of Robinson et al. (1999) for amphiboles from a metamorphosed mafic dyke at Hegenshan, we suggest that the age of c. 245 Ma represents the emplacement time of the Hegenshan ophiolite. Therefore, the final suturing in central Inner Mongolia probably took place in earliest Triassic time (c. 245 Ma). This is reasonably consistent with the general geology of the IMDOM: arc-related granites and volcanic rocks of Permian age on both sides of the Jiaoqier-Xilin Hot belt (Chen et al. 2000; Xiao et al. 2003); the final turbidites of Late Permian age in southern

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Mongolian (Ruzhentsev et al. 1989); Carboniferous shales and siltstones (probably deposited in residual basins) of Late Permian age (Linxin Formation, BGMRIM 1991); and syn- to post-collisional granites with ages of 230 –200 Ma (Chen et al. 2000; Shi et al. 2004). Additionally, the metamorphism related to the final suturing in NE China has recently been constrained at about 240– 215 Ma (Miao et al. 2004; Wu et al. 2004), suggesting suturing broadly simultaneous with that in central Inner Mongolia, although the internal architecture in that sector of the IMDOB still remains poorly understood. Consequently, we suggest that the IMDOB is an Indosinian (Triassic) orogenic belt.

Tectonic evolution Many models have been suggested to delineate the tectonic evolution of the IMDOB (e.g. Wang & Liu 1986; Hsu¨ et al. 1991; Tang & Yan 1993; Sengo¨r & Natal’in 1996; Robinson et al. 1999; Xiao et al. 2003). Most models (e.g. Wang & Liu 1986; Sengo¨r & Natal’in 1996; Xiao et al. 2003) suggested that the tectonic evolution of the IMDOB was characterized by progressive accretion of subduction– accretionary complexes onto opposed continental margins of the North China and Siberian continents, and that the accreted margins collided broadly along the Solon Obo– Linxi belt (the final suture) because the researchers believed that the ophiolites within that zone are the youngest in the IMDOB. This general trend of tectonic evolution is probably universal for most orogenic belts, but the details for the IMDOB may be much more complex. The oldest ophiolite identified in the IMDOB is the Ondor Sum (Tulinkai) ophiolite (.470 Ma) located on the northern side of the North China Craton, and the oldest arc is the Sunid arc (c. 490 Ma) situated on southern margin of the South Mongolia Block. This suggests that the ocean between the two blocks opened and subducted beneath the South Mongolia Block at least during the Cambrian period. In that period, the northern margin of the North China Craton, however, seems to have been a passive continental margin characterized by clastic and carbonate formations. Southerly directed subduction of the ocean beneath the northern margin of the North China Craton was probably initiated during the Late Ordovician–Early Silurian period, as indicated by the Bainaimiao magmatic arc with an age of c. 430 Ma and by the subductionrelated blueschists with Ar– Ar ages of c. 420 Ma (Tang & Yan 1993). The Devonian evolution of the IMDOB is still poorly understood, because geological records are rarely preserved, especially in the northern margin of the North China Craton. It seems that the

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southern margin of the South Mongolia Block during this period was characterized by subduction and accretion to form the Jiaoqier blueschist (c. 380 Ma; Xu et al. 2001) and ophiolite (c. 350 Ma). It is worth noting some points about the Hegenshan ophiolite and the Xilin Gol Complex. According to previous tectonic models, the Hegenshan ophiolite is a component of Early to Mid-Palaeozoic subduction–accretion complexes accreted to the southern margin of the Siberian continent or the South Mongolia Block. However, new dates suggest that the ophiolite is only Late Carboniferous to Early Permian (c. 295 Ma) in age and was probably emplaced in earliest Triassic time (c. 245 Ma), implying that the Hegenshan oceanic basin survived throughout the Permian period. The Xilin Gol Complex, which was previously thought to be a Precambrian micro-continental Block accreted onto the southern margin of the South Mongolia Block, is actually a metamorphosed Palaeozoic accretionary wedge with a lower limit of formation age of 437 + 3 Ma (Shi et al. 2003). Geographically, the Hegenshan ophiolite is located between the Xilin Gol Complex in the south and the Uliastai continental margin in the north. Together with geochemical data for the Hegenshan ophiolite, which are suggestive of a back-arc basin environment for the ophiolite (Robinson et al. 1999), we propose that the Hegenshan ophiolite represents a back-arc basin opened on the Carboniferous margin of the SMB, and that the Xilin Gol Complex and Sunid arc formation are probably a segment separated that from the Carboniferous South Mongolia continental margin during the opening of the Hegenshan basin. This argument is supported by the occurrence of the Baolidao Carboniferous– Permian magmatic arc (310 –275 Ma). Consequently, we conclude that an arc –back-arc basin system, probably resembling the present western Pacific margin, was still active in the southern margin of the South Mongolia Block during the Carboniferous to Permian period. The occurrence of Carboniferous –Permian arc-type volcanic rocks in the Uliastai continental margin and coeval accretionary wedge of Permian age due south of the margin suggests a northdipping subduction of the Hegenshan basin. During this period, the northern margin of North China Block was probably characterized by an Andeantype margin, which was superimposed on the consolidated Early Palaeozoic and/or early Late Palaeozoic accretionary complexes (Xiao et al. 2003). Therefore, the evolution of the IMDOB during the Carboniferous –Permian period is propably a Palaeozoic analogy of the present Pacific. Without reliable geochronological data, the ultramafic –mafic blocks along the northern bank of the Xar Moron River have previously been

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speculated to be the eastern extension of the Early Palaeozoic Ondor Sum ophiolite zone (Wang & Liu 1986; Wu et al. 2002). This speculation was mainly based on the assumption that the Kedanshan ultramafic –mafic complex is an Early Palaeozoic ophiolite, for which only one K –Ar age of c. 570 Ma has been reported. As discussed above, several pieces of evidence demonstrate that the Kedanshan complex, however, is intrusive in origin, and its extremely young age of c. 73 Ma indicates that it is a post-orogenic intrusion. Moreover, other ultramafic–mafic blocks and metamorphic complexes along the northern bank of the Xar Moron River are only late Permian in age, as indicated by the zircon SHRIMP ages of c. 256 Ma and 270 Ma for the Banlashan block and Shuangjing metamorphic complex, respectively. Therefore, it is unreasonable to delineate these blocks and those of Ondor Sum together as a single Early Palaeozoic ophiolite zone. At Ondor Sum, the gabbro (Tulinkai) has an age of c. 460 Ma, whereas the pillow lava (Ondor Sum) has a zircon SHRIMP age of c. 260 Ma, indicating that they are asynchronous. The former does represent remnants of Early Palaeozoic oceanic crust, but its eastern extension is unclear yet, mainly because of either poor exposure of rocks there, or only a local occurrence of the Early Palaeozoic ophiolite. The latter is broadly coeval with those along the northern bank of the Xar Moron River (except for the Kedanshan), and it seems that they may define a Permian ophiolite zone, which occurs immediately north of the Early Palaeozoic zone at present. The ophiolites along the Solon Obo–Linxi belt, or the ‘Solonker suture’ defined by Sengo¨r & Natal’in (1996) and Xiao et al. (2003), have a zircon SHRIMP age of c. 280 Ma (Solon Obo) and a relatively young Rb– Sr isochron age of 262 Ma (Wang & Liu 1986). This indicates that these ophiolites are essentially coeval with those along the Xar Moron River and with that of the Hegenshan ophiolite. This suggests that two oceans or basins coexisted in the IMDOB during the Permian period. The two oceans or basins were separated from each other by a magmatic arc (Baolidao) or a pre-Permian terrane before the final orogenesis of the IMDOB, and were probably closed contemporaneously during the orogenesis. Therefore, we think that no single ophiolite belt in the IMDOB can represent the suture between the Sino-Korean and the Siberian plates; it would be a broad zone comprising at least three of the abovementioned ophiolite belts: the Solon Obo–Linxi, Jiaoqier– Xilinhot, and Erenhot –Hegenshan belts. This would be in accordance with the cryptic nature of the suture zone of the IMDOB. Based on the available data, we show the tectonic evolution of the IMDOB in Figure 8, and

briefly describe it below, with an emphasis on its temporal evolution. In Late Cambrian to Early Ordovician time, probably north-dipping subduction of the Ondorsum oceanic lithosphere (.467 + 13 Ma) beneath the southern margin of the South Mongolia Block (or the Siberian Block) produced the Sunid magmatic arc (c. 490 Ma), whereas the northern margin of the North China Craton was possibly inactive (a passive margin) at about that time (Fig. 8a). In Late Ordovician to Silurian time, the northern margin of the North China Craton evolved into an active margin with the formation of the Bainaimiao continental arc (c. 430 Ma), and the southern margin of the SMB was still active (c. 424 Ma). Accretion of subduction–accretion complexes occurred probably on both the northern margin of the North China Craton and the southern margin of the South Mongolia block (Fig. 8b). In Devonian to Early Carboniferous time, the Hegenshan basin was probably opened on the base of the pre-Permian southern margin of the South Mongolia Block, and a segment of the margin split and drifted southward (present coordinates, Fig. 8c). In the Late Carboniferous to Early Permian period, two oceans or basins (a relict of the Ondorsum ocean and the newly opened Hegenshan ocean) coexisted in the IMDOB; north-dipping subduction of the Hegenshan ocean and two-sided subduction of the Ondorsum ocean probably generated the Uliastai continental arc, the Baolidao magmatic arc (c. 310 Ma) and coeval volcanic rocks (the Dashizhai Formation) in the north, and widespread arc-related Permian volcanic rocks in the south of Linxi (Fig. 8d). In Late Permian time, subduction within the IMDOB had probably ended, with residual basins left, in which the Late Permian carbon-bearing black rock system (the Linxi Formation) was deposited. In Triassic time, final orogenesis of the IMDOB occurred in approximately Early Triassic time (c. 245 Ma), with formation of the Hegenshan granodiorite dykes; syn- and/or post-orogenic magmatism probably continued into Late Triassic time (c. 200 Ma) (Fig. 8e).

Implication for the Mesozoic lithospheric thinning event of the NCC One of the most important tectonic features of the Precambrian North China Craton is the lithospheric thinning event during the Mesozoic, which involved the removal of about 120 km thickness of the lithosphere (Menzies et al. 1993; Zhai et al. 2002),

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Fig. 8. Sketches showing the Phanerozoic tectonic evolution of the Inner Mongolia–Daxinganling Orogenic Belt (IMDOB), China. NCB, North China Block; SMB, South Mongolian Block (see text for details).

although the mechanism of the thinning is still unclear. The wedge-shaped Precambrian North China Craton is bounded by three Phanerozoic orogenic belts: the IMDOB in the north, the Qinling– Dabie –Su-Lu orogenic belts in the south, and the Pacific subduction in the east. As discussed above, the IMDOB is an Indosinian orogenic belt, which is coeval with the Qinling –Dabie –Su-Lu orogenic belt (c. 220 Ma) in the south. At the same time, the Palaeo-Pacific subduction beneath the eastern margin of the North China Craton was probably

continuing. This means that subduction probably occurred simultaneous along the three sides of the NCC during that period. Because the subducted oceanic lithosphere is hydrous, dehydration of this material may trigger intense and extensive partial melting of the lithosphere of the North China Craton. Some of the remnants probably sank because of their high density, and thus caused the thinning of the lithosphere beneath the North China Craton. This may be one of the possible mechanisms for the thinning event.

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This work was financially supported by the National Natural Science Committee of China (grants 40473030 and 40234045) and the Chinese Academy of Sciences (grant KZCX2-104). We thank C. W. Li, H. F. Zhang and Z. Y. Ni for their field support, and H. Tao, Y. R. Shi, J. Ping, Y. H. Zhang, and Z. Q. Yang for their help during the SHRIMP analysis. Discussion with Q. Zhang is appreciated. Two reviewers, W. J. Xiao and B. A. Natal’in, are gratefully thanked for their constructive suggestions.

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Late Mesozoic tectonic transition of the eastern North China Craton: evidence from basin-fill records Z. LI1, Y. LI2, J.-P. ZHENG3 & D. HAN1 1

Institute of Geology and Geophysics, Chinese Academy of Sciences, Beijing 100029, China (e-mail: [email protected]) 2 3

Shengli Oil field, SINOPEC, Dongying 256504, China

China University of Geosciences, Wuhan 420073, China

Abstract: To better understand Mesozoic tectonic transition processes in Eastern China, this paper offers a comparison of Mesozoic basin-fill records around the eastern North China Craton. These Mesozoic basins have similar evolutionary features: inception since the Early Jurassic; basin-fills recording a tectonic evolution from compression and lithospheric thickening before Late Jurassic and/or Early Cretaceous time to intracontinental stretching and lithospheric thinning from Early Cretaceous time; tectonic transition during the late Jurassic with a time lag in shallow crust relative to deep lithosphere. However, basin-fill records reflect two distinct basin systems occurring in the southern and northern margins of the eastern North China Craton. First, varied volcanic rocks including mafic, intermediate –mafic and intermediate –felsic assemblages occur in the Yanshan–Liaoxi basins on the northern margin of the eastern North China Craton from the Early Jurassic to Cretaceous; in contrast, limited calc-alkaline volcanic rocks filled the Hefei basin system on the southern margin of the eastern North China Craton during the Late Jurassic– Early Cretaceous. Second, late Mesozoic lithospheric thinning began at about 163 Ma and 149 Ma in the northern and southern margins, respectively, culminating in basinscale extensional events at about 145 Ma and 132 Ma, respectively. Third, coarse clastic sediments developed in the northern and southern basins during the tectonic transition phase reflect fluvial and alluvial systems, respectively, indicating greater topographic relief in the southern area than in the northern area. Fourth, Mesozoic depocentre migration was complicated in the Yanshan– Liaoxi basin but should a south to north trend in the Hefei basin system. Mesozoic basin-fill depositional and volcanic records in the southern margin of the eastern North China Craton were dominantly controlled by early Mesozoic deep subduction of the Yangtze block and subsequent post-orogenic extension of the Dabie Mountains. On the other hand, basin-fill evolution along the northern margin of the North China Craton was principally controlled by intensive crust–mantle and/or asthenosphere –lithosphere interactions, with regional transition from contractile to extensional strain during Mesozoic time. This study suggests that the late Mesozoic tectonic transition was first induced by crust–mantle interactions in the northern North China Craton and then it extended southwards. Late Mesozoic lithospheric thinning and subsequent tectonic transition are a linked systematic geodynamic process that has no direct relation to the Triassic plate convergence events around the North China Craton.

The tectonic evolution of the North China Craton may be divided into six megastages according to current research (Huang 1960; Wang et al. 1990; Fan et al. 1992; Menzies et al. 1993): (1) Archaean continental nuclei formation; (2) consolidation of the Palaeoproterozoic craton crystalline basement; (3) Mesoproterozoic rifting; (4) Neoproterozoic– Palaeozoic stable cratonic cover deposition; (5) early Mesozoic lithospheric thickening; (6) late Mesozoic–Cenozoic lithospheric thinning. Early Chinese researchers gave more attention to the late Mesozoic tectonic event, termed the ‘Yanshan Movement’, in the North China Craton (Weng 1927, 1929; Zhao 1963). These pioneering works indicated that the ‘Yanshan Movement’ marks a tectonic regime transition of the North China Craton

from the Tethyan to the Circum-Pacific continental margin tectonic domain (Huang 1960; Wang et al. 1990). Since the 1990s, with the development of continental geodynamic research, geoscientists have renewed their interest in the Yanshan–Liaoxi and Dabie Mountains–Sulu belts around the North China Craton. By virtue of new ideas and technologies, a new round of fundamental geological research has begun, and many geodynamic models of the late Mesozoic tectonic transition in East China have been discussed and published (Fan et al. 1992; Menzies et al. 1993; Chen 1998; Davis et al. 1998, 2001; Deng et al. 2000, 2004; Zheng et al. 2000; Zhai et al. 2004; Zhao et al. 2004). Basin-fill deposits contain much information on tectonic, volcanic and climate evolution, and their

From: ZHAI , M.-G., WINDLEY , B. F., KUSKY , T. M. & MENG , Q. R. (eds) Mesozoic Sub-Continental Lithospheric Thinning Under Eastern Asia. Geological Society, London, Special Publications, 280, 239–266. DOI: 10.1144/SP280.12 0305-8719/07/$15 # The Geological Society of London 2007.

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analysis is very important for understanding the tectonic framework and lithosphere geodynamic processes (e.g. Dickinson & Suczek 1979; Burbank & Reynolds 1988; Graham et al. 1993; Brozovic & Burbank 2000; Hendrix 2000; Kuhlemann et al. 2006; Najman 2006). Compared with research on structures, magmatism and metallogenesis, sedimentary research on the Mesozoic basins along the northern and southern margins of eastern North China has lagged behind. Even systemic knowledge regarding regional depositional evolution at basin scale has been lacking. Some published integrated basin research in adjacent areas, e.g. Mongolia (Graham et al. 2001; Hendrix et al. 2001; Johnson et al. 2001; Badarch et al. 2002; Johnson 2004) and NE Asia (Ren et al. 2002; Meng et al. 2003) principally discussed extensional processes of late Mesozoic age rather than tectonic transition from early to late Mesozoic time. However, some tectonic problems still exist between Mongolia and NE China, as presented by Graham et al. (2001). Recently, new opportunities for improving this situation have emerged because of the publication of many isotopic geochronological, stratigraphic and sedimentary data for North China. This paper mainly offers a comparison of basin-fill sequences, depositional systems and their regional changes in Mesozoic (rather than late Mesozoic) basins on the northern and southern margins of the eastern North China Craton, and affords insight into geodynamic mechanisms and implications for the late Mesozoic tectonic transition based on an integrated discussion of sedimentary, structural and volcanic data.

Regional tectonic setting and geological framework of the Mesozoic basins The eastern North China Craton is situated between the late Palaeozoic (probably to Triassic) Mongol accretionary orogen (Wang et al. 1990; Enkin et al. 1992; Graham et al. 2001; Badarch et al. 2002) to the north and the Triassic Dabieshan– Sulu collisional orogen (Okay et al. 1993; Maruyama et al. 1994) to the south (Fig. 1). The Late Mesozoic tectonic transition of the North China Craton, from lithospheric thickening to thinning, fostered a series of Mesozoic basins around the craton. The Yanshan –Liaoxi region, in the eastern segment of the east –west-striking Yinshan–Yanshan belt (Zheng et al. 2000), is situated along the northern margin of the North China Craton (Fig. 1). Because a collision between the North China Craton and the Mongolian orogen occurred along the Solonker (Silamulun or Suolun) suture in the latest Palaeozoic–earliest Mesozoic, the Jurassic –Cretaceous Yinshan–

Yanshan orogen is considered to be the result of intracontinental tectonism. The geological framework of the Yanshan–Liaoxi region is mainly characterized by early east –west structures superimposed by late NNE–SSW structures and cross-cut by structural units that include metamorphic core complexes and magmatic diapirs (Song 1999). Because of late intensive intracontinental deformation, most Mesozoic basins are incompletely preserved in the Yanshan–Liaoxi area. According to available stratigraphic records, the Mesozoic remnant basin systems consist of north–south belts, separated by the east –west-striking Shanyi– Pingquan fault (Fig. 2a). West and east subareas in the northern belt are named Zhangbei– Guyuan and Fengning–Weichang, respectively (Fig. 3). Stratigraphic sequences in these subareas began in the Early Cretaceous and ended in the Late Cretaceous. The stratigraphic subareas in the southern belt are named Beijing– Chengde and Liaoxi, respectively, from west to east (Fig. 3). These subareas commonly are characterized by more complete stratigraphic sequences from Triassic to Early Cretaceous, and locally even Late Cretaceous. Within the southern belt of the Yanshan– Liaoxi area mentioned above, western basins include (from north to south) Zhangjiakou, Xuanhua, Huailai and West Beijing. These basins are separated by fold– thrust structures. Central–eastern basins mainly include Luanping –Chengde, Kuancheng, Qianxi and Jianchang, whose boundary faults show twostage activity indicated by mylonites and breccias resulting from early thrusting and late extension. On the other hand, Jurassic and Cretaceous basins in the Yanshan–Liaoxi region show different structural styles. The Jurassic basins are mainly controlled by nearly east –west structures with thrust boundaries. However, the Cretaceous basins are distributed across the southern and northern belts of Yanshan–Liaoxi, whose boundary faults mostly strike NE–SW or NNE–SSW and have extensional features. As the boundary between the North China and Yangtze Cratons, the Dabie Mountains orogenic belt mainly resulted from early Mesozoic collisional and late Mesozoic post-collisional orogenesis (Okay et al. 1993; Xu et al. 1994, p. 135–137; Li, Z. et al. 2004). The Hefei basin system, striking nearly east –west, is located in the northern foothills of the Dabie Mountains (Fig. 2b), which are themselves bounded to the south, east and north by the Xiaotian –Mozitan, Tanlu and Yingshang – Dingyuan faults, respectively. The Hefei basin system may also be divided into north and south belts, because thicknesses, sedimentary lithofacies, burial history and thermal history of the Mesozoic successions are mutually different.

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Fig. 1. Schematic map showing tectonic units of North China and adjacent areas (modified from Enkin et al. 1992).

The southern belt of the Hefei basin system, termed the ‘Beihuaiyang fold belt’, is an elongate, nearly east –west zone, which is bounded by the Liuan –Qieshan fault to the north and the Xiaotian– Mozitan fault to the south. This belt is 400 km in length and 20 –40 km in width, with a basement consisting of greenschist-facies metamorphic rocks, termed the Foziling and Meishan groups of Neoproterozoic –Palaeozoicage, and is filled by Jurassic –Cretaceous clastic rocks interbedded with volcanic and volcaniclastic rocks. The Beihuaiyang belt lies between the typical North China Craton and Dabie Mountains orogenic belt and has a complicated evolution (Yang 1982; Xu et al. 1994, p. 13–19; Fan et al. 2004). The northern belt of the Hefei basin system is bounded by the Liuan– Queshan fault to the south (Fig. 2b). Its basement mainly consists of migmatitic metamorphic rocks of the Huoqiu and Wuhe groups, and shows distinct tectonic evolutionary differences from the southern belt. The Jurassic– Cretaceous strata of the northern belt mostly can be correlated with strata of the southern belt (Li, Z., et al. 2001; Han 1996), but their lithofacies and thicknesses differ.

Remarkably, a single remnant basin of Late Jurassic –Early Cretaceous age, the Xiaotian basin, occurs at the southern edge of the Hefei basin system (Fig. 2b). In this basin, Late Jurassic– Early Cretaceous intermediate –felsic volcanic rocks directly overlie the metamorphic basement of the Foziling group and the basin was evidently controlled by the north-dipping Xiaotian –Mozitan fault.

Basin-fill sequences Yanshan– Liaoxi area, northern margin of the North China Craton Varied clastic and volcanic rock assemblages occur across the Yanshan–Liaoxi area, along the northern margin of the North China Craton. Stratigraphic and petrological research reveals correlated basin-fill sequences and unconformities in Mesozoic basins, indicating similar tectonic and climatic controls. In spite of structural segmentation and deformation in the study area, six basin prototypes can be distinguished from these remnant basins, which reflect

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Fig. 2. Structural framework and Mesozoic stratigraphic outcrops in the Yanshan– Liaoxi area (a) and the south Hefei basin system (b), located in the northern and southern margins of the eastern North China Craton as shown in Figure 1. Sedimentary profiles observed are marked by black triangles. 1, Shangyi–Pingquan fault; 2, Zhangjiakou fault; I, Jinzhai– Huoshan–Shucheng fault; II, Xiaotian–Mozitan fault.

Mesozoic tectonic processes from phase I to phase VI (Fig. 3). Jurassic basins in phases I, II and III. The phase I basins in the Yanshan–Liaoxi area, represented by the Lower Jurassic Xinshikou Fm. (J1x), are mainly preserved in the central–eastern part of the study area, with an unconformity separating it from

Lower Triassic rocks (Chen & Wu 1997). The Xinshikou Fm., 20–610 m thick, shows an upwardcoarsening sequence. The unit contains extensive orthoconglomerates mostly consisting of subrounded gravels with a diameter of 5–20 cm (Fig. 4a) and complicated compositional assemblages (Fig. 5a and b), attributed to a gravelly braided river depositional system (Li, Z. et al. 2004). We infer that the

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Fig. 3. Mesozoic chronostratigraphic correlation across the southern and northern margins of the eastern North China Craton.

depositional landform of the Xinshikou Fm. was characterized by larger catchment areas and a small difference of elevation between basins and mountains. Little debris-flow deposition in this phase also may be related to a wet environment, as indicated by Parrish (1993) and Zhou (1995). The phase II basins were first filled by the Nandaling Fm. (J1n), in unconformable contact with the underlying Xinshikou Fm., but a disconformity is overlain by the Xiahuayuan Fm. (J2x) and the Jiulongshan Fm. (J2j), indicating a new basin generation. The Lower Jurassic Nandaling Fm. occurs only in the southern zone of the Yanshan –Liaoxi area and mainly consists of basalt and basalt andesite 300– 600 m thick. The Nandaling Fm. reflects a local extensional event after early Jurassic compression (Li, X., et al. 2004). The Xiahuayuan Fm., distributed south of the Shanyi–Pingquan fault and covering a larger depositional area, mainly consists of grey–green, grey–yellow siltstones, dark grey mudstones, shales and fine sandstones interlayered with thin coal seams, coarse sandstones and marl with a total thickness of 500 –800 m, principally showing an evolutionary sequence from undercompensated lacustrine to fluvial–delta-dominating depositional system

(Fig. 5c and d). The Jiulongshan Fm., whose depositional area is larger than that of the Xiahuayuan Fm., disconformably overlies underlying strata and shows multiple cycles of gravelly sandstones and conglomerates to brown mudstones. The Jiulongshan Fm. is characterized by a depositional sequence with a coarse–fine –coarse grain-size change (Fig. 5c and d). Fan-delta deposition, with typical upward-coarsening sequences from thin sandy distributary channels to thick gravelly braided channels (Fig. 4b and c), is developed in both lower and upper segments of the Jiulongshan Fm. Thick brown mudstones generally interlayered with thin conglomerates and sandstones with normal graded bedding developed in the middle segment, most probably by gravity flow related to a fan-delta front environment. In addition, fluvial lag deposits, comprising angular lava clasts mostly less than 15 cm in diameter, occurred at the bottom of the Jiulongshan Fm. The abundance of basalt andesite and tuff gravel clasts, with diameters less than 8 cm, increases upwards in the Jiulongshan Fm., indicating an increasing influence of volcanic provenances. The phase III basins in the Yanshan–Liaoxi area developed since deposition of the Tiaojishan Fm.

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Fig. 4. Stratigraphic and sedimentary observations in the field of the Yanshan–Liaoxi region: (a) gravelly braided river deposition of the Xinshikou Fm. in the Dashipeng section, Luanping; (b) fan-delta deposition of the lower Jiulongshan Fm. in the Dashipeng section, Luanping; (c) a typical upward-coarsening sequence of the lower Jiulongshan Fm. in the Dashipeng section, Luanping; (d) gravelly braided river dominating deposition of the Tuchengzi Fm. in the Hiying section, Luanping; (e) volcanic rocks dominating gravel assemblages of the Tuchengzi Fm. in the Changshanyu section, Luanping; (f) limestones dominating gravel assemblages of the Tuchengzi Fm. in the Dazhangzi section, Kuancheng; (g) Zhangjiakou Fm. resting unconformably on Tiaojishan Fm., Chengde; (h) graded bedding in the Liying section, Luanping.

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Fig. 5. Detrital compositional assemblages of conglomerate and sandstone and depositional sequences in the Jurassic profiles of Yanshan–Liaoxi: (a) Xinshikou Fm. in the Shanggu section, Chengde; (b) Xinshikou Fm. in the Dashipeng section of the Luanping basin; (c) Xiahuayuan Fm. and Jiulongshan Fm. in the West Beijing basin; (d) Xiahuayuan Fm. and Jiulongshan Fm. in the Dashipeng section of the Luanping basin. All section locations are shown in Figure 2a. Gravel compositional symbols and contents: 1, gneiss (%); 2, quartzite (%); 3, sillicolites (%); 4, clastic rock (%); 5, carbonate (%); 6, granite (%); 7, lava (%).

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volcanic rocks (J3tj). The Tiaojishan Fm., 1000– 3400 m thick, is composed of calc-alkaline trachyandesites with a high potassium content and results from extensive volcanic activity principally during the end of the Middle Jurassic at 160.7 – 163 Ma (Davis et al., 2001; Liu et al. 2003). Zhang et al. (2001) suggested that the Tiaojishan Fm. lava is adakitic, resulting from thickened crust. They also indicated that the Tiaojishan Fm. lava was derived from an enriched mantle source and resulted from partial melting of old (Archaean) basaltic rocks of the crust –mantle transition zone through magmatic underplating in an intracontinental compressive setting. The Tuchengzi Fm. (J3t), with a thickness of 800– 1400 m, disconformably overlies the Tiaojishan Fm. and shows a three-phase evolution from alluvial fan– gravelly braided river, through fan-delta –lacustrine, to gravelly braided river depositional systems in the Luanping– Chengde sections (Fig. 6a and b), with a dominant gravel composition changing upwards from andesite to rhyolite. Gravelly braided river deposition is principally characterized by vertical accretion of thick conglomerates with large-scale trough beddings and scour pits (Fig. 4d), and fan-delta–lacustrine deposition is characterized by thick dark brown mudstones and interlayered thin sandy conglomerates and sandstones with normal graded bedding or scour structures. Coarse clastic deposits occur widely in the Tuchengzi Fm. and show varied gravel compositional assemblages in different profiles over the Yanshan –Liaoxi area (Fig. 4e and f), indicating distinct provenances between the southern and northern sides of the Shangyi– Pingquan fault. Cretaceous basins in phases IV, V and VI. The phase IV basins developed after the Zhangjiakou Fm. (K1z) and Dabeigou Fm. (K1d) volcanism of Early Cretaceous age and mostly occur in the central–western part of the northern belt of Yanshan –Liaoxi. The Zhangjiakou Fm. (K1z), with ages in the range 143.9 –136.3 Ma (Li, W. et al. 2001; Liu et al. 2003; Niu et al. 2003), is composed of rhyolites and trachytes with minor mafic lavas and interlayered volcaniclastic rocks with a thickness of 622– 3867 m. The unit rests unconformably on underlying strata (Fig. 4g). Geochemical data indicate that the Zhangjiakou Fm. resulted from partial melting of the lower crust as a result of intensive magmatic underplating with an evident mantle source influence (Shao et al. 2003). Above the Zhangjiakou Fm., the Jiufotang Fm. (K1j) mostly occurs in the central–western part of the Yanshan –Liaoxi area. The Yixian Fm. (K1y) occurs further east. These lithostratigraphic units principally consist of a series of lacustrine and fandeltaic deposits with a total cumulative thickness of

1500–2400 m, and are generally characterized by interstratified volcanic rocks, dark grey mudstone– carbonaceous shales and thin coal seams (Fig. 6c). In particular, the Jiufotang Fm. (K1j) frequently shows multi-cyclic deposits of mudstones, siltstones and gravel-bearing sandstones, in which most gravel-bearing sandstones with graded bedding are distributed along normal faults and indicate gravity-flow deposition (Fig. 4h). The coeval Yixian Fm. (K1y) is characterized by lacustrine depositional sequences interlayered with intermediate to mafic volcanic rocks with a high potassium content and calc-alkaline compositions, which reflect an intracontinental stretching environment (Li, W., et al. 2002). The phase V basins mainly occur in the western and eastern margins of Yanshan–Liaoxi. The western basins consist of the Xiadian Fm. (K1x) and Qingshila Fm. (K1q), 600– 800 m in total thickness, and dominantly are filled by grey gravelbearing sandstones, dark grey mudstones and shales interlayered with thin coals, reflecting an intracontinental lacustrine environment (Chen & Wu 1997). At the same time, the eastern basins are filled by the Shahai Fm. (K1s) and Fuxin Fm. (K1f), which are approximately 2000 m thick. These basin-fills have typical lacustrine –fan-delta depositional sequences with more frequent gravityflow deposition than in the western basins, which are considered as fault basins that developed in an extensional setting (Li 1988). The phase VI basins, similar to those of phase V, are also distributed in the western and eastern margins of Yanshan–Liaoxi. The western and eastern basins are filled by the Nantianmen Fm. (K2n) and Sunjiawan Fm. (K2s), respectively, in unconformable contact with underlying older strata. The Nantianmen Fm., 500–900 m thick, is composed of yellow –brown and grey– white conglomerate, sandy conglomerate and interlayered mudstone, attributed to braided river and alluvial plain deposition (Chen & Wu 1997; Li, Z., et al. 2004). Quartz trachyte –tuff gravel assemblages commonly occur near the base of the Nantianmen Fm., and granite –quartzite –quartz porphyry–trachyte gravel assemblages at the top, reflecting an intensive unroofing of the adjacent provenance area. Most clasts are subrounded to rounded with average diameters in the range 5–10 cm. The Sunjiawan Fm., occurring in the Liaoxi area, has similar depositional facies and sequences to the Nantianmen Fm. (after Li 1988).

Hefei basin system, southern margin of the North China Craton The Jurassic–Cretaceous sequence of the southern zone of the Hefei basin system, 5000–7000 m

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Fig. 6. Late Jurassic–Early Cretaceous depositional sequences and detrital compositional assemblages in the Yanshan– Liaoxi region: (a) Tuchengzi Fm. in the Changshanyu section of the Luanping basin; (b) Tuchengzi Fm. in the Dazhangzi section of the Kuancheng basin; (c) Jiufotang Fm. in the Liying section of the Luanping basin. All section locations are shown in Figure 2a. Gravel compositional symbols and contents: 1, gneiss (%); 2, quartzite (%); 3, sillicolites (%); 4, clastic rock (%); 5, carbonate (%); 6, granite (%); 7, lava (%).

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thick, is well exposed in the Huoshan–Jinzhai region, and is composed of terrigenous clastic deposits interbedded with volcanic and volcaniclastic rocks. On the other hand, Jurassic strata of the central–northern zone of the Hefei basin system are well exposed in the Feixi region, where they are 4000–5500 m thick, and, in general, consist of a suite of terrigenous clastic deposits with few volcaniclastic rocks. Six phases of basin development can be distinguished from the Mesozoic basin-fill sequence of the Hefei basin (Fig. 3). Jurassic–Early Cretaceous basins in phases I, II and III. The phase I basins along the southern margin of the North China Craton occur only in the central–northern zone of the Hefei basin system and are filled by the Fanghushan Fm. (J1f), Fm. About 400 m thick, which unconformably overlies grey–black graphite schists and grey mica–quartz schists of the pre-Mesozoic Foziling Group and consists of three sedimentary genetic units (Fig. 8c). For lower genetic units of the studied profiles, 30–40 m thick, grey to brown – grey sandy conglomerates and orthoconglomerates are developed, in which the gravel clast sizes are generally less than 0.2 m. Sorting is good, and clasts are rounded and locally imbricated, reflecting deposition from sandy–gravelly meandering streams (Li, Z. et al. 2001). Grey muddy siltstones, interbedded with thin-layered shales and coal seams, occur in the middle units, indicating floodplain deposition (Han 1996). Finally, the upper units are composed of greyish white gravel-bearing sandstones interbedded with thin layers of siltstones–fine sandstones and conglomerates, in which sandstones have high compositionally maturity with large-scale cross-bedding, reflecting sand-bar deposition from meandering streams. Palaeocurrent indicator directions reflect flow to the NE and east (Li, Z. et al. 2001). In the southern Hefei basin, the phase II basins were filled by the Sanjianpu Fm. (J2s), more than 1700 m thick (Fig. 8a and b), which consists of grey and brown silty mudstones and siltstones interstratified with fine conglomerates with ubiquitous wedge-shaped inclined bedding (Fig. 7a) and large-scale trough cross-bedding (Fig. 7b), indicating gravel –sand meandering and braided streams. Dominant palaeocurrent directions are to the NE (Li, Z. et al. 2001). At the same time, the Yuantongshan Fm. (J2y), about 1000–3000 m thick, was developed in the northern Hefei basin. This unit is composed of purple–red siltstone, silty mudstone and, greyish green silty mudstone interbedded with gravel-containing medium–fine sandstone (Fig. 8c). Climbing ripples, bioturbation, burrows and interference ripples (Fig. 7c and d) reflect

typical fine-grained deposition from meandering streams and flood plain (Miall 1996, p. 191–250). The phase III basins were filled by upper Jurassic and lower Cretaceous stata that include the Fenghuangtai Fm. and the Zhougongshan Fm. in the southern and northern zones of the Hefei basin, respectively. The Fenghuangtai Fm., more than 1800 m thick, disconformably rests upon the underlying Sanjianpu Fm. and includes three upwardcoarsening sedimentary–genetic units (Fig. 8a and b) mainly consisting of mottled and grey–purple polymictic conglomerate (Fig. 7e), in which lenticular sandstone bodies with large-scale inclined bedding and scour surfaces are common (Fig. 7f), suggesting typical alluvial fan deposition dominated by debris flows. In contrast, the Zhougongshan Fm. is about 900–1300 m thick and is composed of purple– red gravel-bearing coarse sandstone and grey sandstone (Fig. 8c). Well-developed sedimentary structures include scour marks and trough cross-bedding (Fig. 7g), suggesting sedimentary environments of a gravelly– sandy braided river and alluvial plain. The Maotanchang Formation dominantly occurs in the southern Hefei basin and is conformable with or is in facies relations with the Fenghuangtai Fm. (Li, S. et al. 2002). This unit consists of three genetic –stratigraphic units with a total thickness about 900– 1000 m. The lower segment shows six cycles of intermediate –acidic volcanic eruption and consists of andesite, andesite tuff, and volcaniclastic breccias. The middle segment is composed of grey silt-bearing mudstones in thin layers, with the fossil Estheria reflecting lacustrine facies (Fig. 7h). The upper segment shows three cycles of intermediate –acidic volcanic eruptions that consist of andesite, trachytic crystalline clastic tuff with ages of 149–138 Ma (Wang et al. 2002). Cretaceous basins in phases IV, V and VI. The phase IV basins were filled by the Heishidu Fm. (K1h) and Zhuxiang Fm. (K1z) in the southern and northern zones of the Hefei basin, respectively. The Heishidu Fm. is about 110–650 m thick and comprises two genetic units. The lower segment consists of mottled sandy, conglomerates and volcanic breccias, with blocky structure and scour surfaces, indicating alluvial fan or fan-delta deposition (Fig. 9a and b). In addition, thin intermediate –mafic lavas mainly including basaltic – trachytic andesites are interbedded in the lower segment and are dated at 132– 116 Ma (Wang et al. 2002). The upper segment shows thick dark grey mudstone interbedded with gravel-bearing sandstone and fine sandstone with slump structures and graded beds, probably indicating typical lacustrine and fan-delta deposition. The Zhuxiang Fm. is about 1000 m thick and consists of interbedded dark brown or brownish grey silty mudstones and

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Fig. 7. Stratigraphic and sedimentary observations in the field of the Hefei basin: (a) wedge-shaped inclined bedding of the Sanjianpu Fm., Shucheng; (b) large-scale trough cross-bedding of the Sanjianpu Fm., Huoshan; (c) climbing ripples of the Yuantongshan Fm., Feixi; (d) interference ripples in brown siltstone mudstones of the Yuantongshan Fm., Feixi; (e) polymictic conglomerate of the Fenghuantai Fm., Huoshan; (f) scour marks of the Fenghuantai Fm., Huoshan; (g) trough cross-beds of the Zhougongshan Fm., Feixi; (h) mudstones of a thin layer of the Maotanchang Fm., Xiaotian.

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Fig. 8. Jurassic– Early Cretaceous depositional sequences in the Hefei basin: (a) Sanjianpu Fm. and Fenghuangtai Fm. in the Fenghuangtai profile, Huoshan; (b) Sanjianpu Fm. and Fenghuangtai Fm. in the Dushan profile, Liuan; (c) Fanghushan, Yuantongshan and Zhougongshan Fms. in the Fanghushan profile, Feixi. All section locations are shown in Figure 2b.

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Fig. 9. Cretaceous depositional sequence and detrital compositional assemblages in the Hefei basin: (a) Heishidu Fm. and Xiaotian Fm. in the Xiaotian profile, Shucheng; (b) Heishidu Fm. in the Huoshan profile; (c) Zhuxiang Fm. and Xiangdaopu Fm. from well data in the northern Hefei basin; (d) Zhangqiao Fm. from well data in the northern Hefei basin. All section locations are shown in Figure 2b.

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volcaniclastic fine sandstones, with brownish red conglomerate and gravel-bearing coarse sandstone developed at the bottom, indicating deposition of sandy delta –lacustrine systems (Fig. 9c). The phase V basins were filled by the Xiaotian Fm. (K1xt) and Xiangdaopu Fm. (K1xd) in the southern and northern Hefei basin, respectively. The Xiaotian Fm. is greater than 800 m thick and mainly comprises thick dark grey muddy shale with horizontal bedding, interbedded with thinlayered siltstone and fine sandstone with horizontal beddings and small-scale cross-beds, indicating typical lacustrine facies (Fig. 9a). The Xiaotian Fm. overlies Tertiary purple–red and thick sandy conglomerates with an angular unconformity in the southern Hefei basin. The Xiangdaopu Fm., 800– 900 m thick, dominantly shows interbedded dark greyish and grey brown silty mudstone, fine sandstone and siltstone layers, with common gypsum laminae and horizontal bedding (Fig. 9c), and reflects a restricted lacustrine environment (Han 1996; Li, Z. et al. 2001). The phase VI basins occur only in the northern Hefei basin and are filled by the Zhangqiao Fm., which is 900–1000 m thick. This unit consists of two sedimentary–genetic units: a lower interbedded dark brown fine sandstone and silty mudstone, with current bedding, and an upper brownish red, greyish purple thick-layered conglomerate and sandstone unit interbedded with thin silty mudstone (Fig. 9d). Trough cross-beds and scour marks are common at the bottom, probably indicating sandy braided river and alluvial plain deposition (Li, Z. et al. 2001). Briefly, the basin-fills of the Yanshan –Liaoxi area are characterized by two second-order sequences (Miall 1996, p. 75 –98). The lower second-order sequence consists of upward-thinning and upward-coarsening depositional cycles of Early to Late Jurassic age, interlayered with two series of intermediate –mafic volcanic rocks. The upper second-order sequence principally shows an upward-thinning and upward-coarsening depositional cycle of Early to Late Cretaceous age, in which, notably, volcanic rocks are developed only in the beds of Early Cretaceous age, and are attributed to intermediate –felsic eruptions. This basin-fill probably indicates an evolutionary process with basin structural-type changes. As shown in Figure 10, this process includes a compressional cycle, from phase I to pahse III, and a rifting cycle from initial rift (phase IV), through intensive rift (phase V), to post-rift (phase VI). The Hefei basin also shows basically similar basin-fill sequences to the Yanshan –Liaoxi area with distinct volcanic interlayers; that is, an upward-thinning and upwardcoarsening depositional cycle of Jurassic –Early Cretaceous age, exists, with volcanic rocks only at

the top, and an upper second-order depositional cycle of Early to Late Cretaceous age, that is similar to that in the Yanshan–Liaoxi area, with early developed volcanic rocks (Fig. 10).

Depositional framework and basin migration Based on the basin-fill sequences and depositional facies features described above, and the sedimentary system distribution, an integrated palaeotectonic and palaeogeographical reconstruction provides some important information on basin formation, distribution and migration in the southern and northern margins of the eastern North China Craton.

Yanshan – Liaoxi area, northern margin of the North China Craton Early Jurassic. Depositional assemblages and their distribution indicate that the Early Jurassic relief of central Yanshan is characterized by a north to south topographic gradient. In the southern lowland, depositional basins mainly occurred in two similarly east– west-striking belts that probably resulted from flexural loading (He et al. 1999). The northern basin belt was approximately along the Shanyi– Pingquan fault and was filled by gravelly alluvial fans and gravelly braided river depositional systems. Grain-size distributions fine from north to south, and the average palaeocurrent direction obtained from 133 measurements is 173 + 158, based on a weighted mean calculation (Li, Z. et al. 2004). We interpret these observations to indicate that the main provenance was from the northern highland. The southern basin belt was approximately along the strike of the current West Beijing syncline and was filled by gravelly braided river depositional systems, with relatively finer overall grain-size distributions than that of the northern basin belt. We infer that no deposition occurred north of the Shanyi–Pingquan fault. Middle Jurassic. As shown in Figure 11a, we infer that the lower Middle Jurassic Xiahuayuan Fm. subsequently was deposited in the Shanyi, Xuanhua, Chicheng, West Beijing, Chengde, Jianchang, Beipiao sub-basins, in a broad east–west-striking depositional facies zone parallel to the pre-existing palaeotectonic strike. Depositional assemblages and sequences of the Xiahuayuan Fm. indicate lacustrine environments, passing upward to coeval delta and meandering river environments from basin margin to centre. This also shows that locally persistent spreading towards the south and the north occurred in the southern belt of

Fig. 10. Basin-fill sequences and their evolutionary phases in the southern and northern margins of the eastern North China Craton.

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Fig. 11. Depositional systems and regional provenance reconstruction of the remnant basins in the Yanshan–Liaoxi area; compiled and revised from Li, Z. et al. (2004), Qi et al. (2003) and Liu et al. (2004). (a) Middle Jurassic; (b) Middle-Late Jurassic; (c) Early Cretaceous.

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Yanshan –Liaoxi. Some small coal-bearing basins are also found along the Shanyi– Pingquan fault belt, probably reflecting an inactive tectonic setting. Finally, at the end of the Middle Jurassic, the Jiulongshan Fm. was deposited at a time of gradual contraction of the lacustrine area. In the middle segment of this unit, fan-delta deposits probably reflect an ephemeral extensional environment. Evidently, the whole structural framework of the Yanshan area still retained its overall east –west strike through the Early– Middle Jurassic. Middle –Late Jurassic. The Tiaojishan Fm. is mainly composed of andesite, trachyte and rhyolite that unconformably overlie the Jiulongshan Fm. in the southern Yanshan and the Archaean metamorphic basement in the northern Yanshan. A similar stratigraphic distribution exists between the Tiaojishan Fm. and Tuchengzi Fm., with differing styles of basin-fill sequences and framework from those of pre-existing basins (Fig. 11a and b). According to our palaeogeographical reconstruction of the Tuchengzi Fm., most basins of this phase were distributed along the Longhua, Shanyi–Pingquan, Xuanhua–Xiahuayuan and Jianchang–Chaoyang thrust belts. Fluvial conglomerates were developed proximal to the thrust fronts and relatively fine delta–lacustrine systems occurred in more distal regions (Fig. 11b). Therefore compared with the preexisting Middle Jurassic basins, thrust structures probably exerted an important control on Middle– Late Jurassic basin formation. The leading edge of these thrusts evidently migrated northward. Furthermore, Tuchengzi Fm. deposition during this time shows new sub-basin development and structure segmentation along a NE strike. Widespread thick coarse clastic deposition, with varied provenance assemblages, indicates an evident sub-basin segmentation and an increase in sub-basin subsidence rates at this time. According to 455 palaeocurrent measurements of the Upper Jurassic Tuchengzi Fm., palaeocurrent directions dominantly trend 161 + 158 and 121 + 158 in weighted mean in the central and western domains, respectively, of the southern zone of Yanshan –Liaoxi (Li, Z. et al. 2004). These observation indicate major provenances from the NW (Fig. 11b). The basin distribution of this phase is characterized by a single basin with a NE strike, and basin groups or zones with an east– west strike, and a similar basin system persisted until the Early Cretaceous. This overall basin structure was identified earlier from the Duolun basin group north of Yanshan – Liaoxi (Li 1988). Early Cretaceous. Compared with the Tiaojishan Fm., which is mostly distributed south of the Shanyi–Pingquan fault, the Zhangjiakou and

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Dabeigou Fms. mainly occur north of the Shanyi– Pingquan fault and west of the Luanping– Weichang, and are evidently controlled by NNE –SSW-striking basin structures. These lithostratigraphic units show evidence of thickening with increasing age from east to west, probably resulting from volcanic activity migrating and strengthening from east to west. In fact, from the regional stratigraphic thickness map of NE Asia, the Early Cretaceous volcanic rocks were also controlled by structures with a NNE–SSW strike, with the thickest domain west of Daxinganling, NE China (Li 1988). After the early Cretaceous volcanic activity, fluvio-lacustrine deposition developed over the Yanshan– Liaoxi area. Major depositional depocentres of this phase show both east and west migrating trends. As shown in Figures 3 and 11c, the synchronous Yixian and Jiufotang Fms. were dominantly developed in Liaoxi and western Yanshan, respectively, far from central Luanping– Weichang. Subsequently, the late Early Cretaceous, Xiadian–Qingshila Fm., was mainly deposited in the Fengning–Wanquan domain in the, western part of Yanshan. At the same time, the Shahai–Fuxin Fm. was locally deposited in the Liaoxi domain. Little depositional record of late Early Cretaceous age can be found in the central Luanping–Weichang domain. Evidently, similar basin distributions existed through early and late Early Cretaceous time. Furthermore, the western Yanshan basins appear to have been dominantly controlled by basinmarginal normal faults with a westerly dip. However, the eastern Yanshan basins are controlled by basin-marginal normal faults with an easterly dip. This structural configuration suggests a symmetrical basin-and-range structure relative to the central Luanping– Weichang belt during the Early Cretaceous. These Early Cretaceous structural styles are similar to those of Late Mesozoic extensional basins of NE China and Mongolia, the tectonic drivers of which remain uncertain and include large-scale tectonic stress transition (Li 1988; Ren et al. 2002), metamorphic core complexes and/or collapse of bounding orogenic belts (Davis et al. 2001, 2002; Graham et al. 2001; Johnson et al. 2001; Meng et al. 2003). Late Crateceous. Palaeogeographical analysis shows that the coeval Sunjiawan and Nantianmen Fms., dominantly composed of conglomerate and sandy conglomerate deposited from braided rivers and alluvial fans, mostly occur in eastern and western Yanshan (Chen & Wu 1997; Li 1988), respectively, evidently indicating depositional relief similar to that of the Early Cretaceous (Fig. 11c). In other words, the Late Cretaceous deposition was still controlled by extensional

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structures with a NNE–SSW strike and with relative uplift along the Luanping –Weichang belt, central Yanshan. This uplift belt in central Yanshan probably extends southward to the Cangxian structure belt, a Late Cretaceous –Palaeogene uplift with a NNE–SSW strike in the Bohai basin, North China (Allen et al. 1997).

Hefei basin, southern margin of the North China Craton The fill-sequences of the Hefei basin record significantly different lithofacies and depositional architectures in Lower Jurassic, Middle Jurassic, Upper Jurassic –Lower Cretaceous and Upper Cretaceous strata. These differences, as shown in Figure 9, probably reflect changes in geodynamic processes. Early Jurassic. Meandering stream– fluvial plain systems developed in the Hefei basin during Early Jurassic time. Dominant palaeocurrents mainly trend towards the NE in the southern Hefei basin and to the NE or east in the central Hefei basin (Li, Z. et al. 2001), reflecting the development of axial as well as transverse fluvial systems. Stratigraphic thickness maps also indicate that depocentres of the Early Jurassic are restricted to the southern margin of the Hefei basin (Figs 12 and 13a). Therefore the Early Jurassic basin probably reflects the development of a piedmont, north of the Dabie Mountains, from weak flexure of the lithosphere with a relatively small topographic height difference (Fig. 14a). Middle Jurassic and Late Jurassic –Early Cretaceous. Seismic profiles and stratigraphic thickness maps show that the main depocentres of the Middle Jurassic and Late Jurassic –Early Cretaceous strata, including the Sanjianpu and Fenghuangtai Fms. in the south and the Yuantongshan and Zhougongshan Fms. in the north, are also

restricted to the southern margin of the Hefei basin (Figs 12 and 13b). Evidently, deposition of this phase was controlled by the Jinzhai – Huoshan– Shucheng fault with a southward dip (Fig. 12), probably implying that basin-fill to the north resulted from a flexural setting with marginal thrusting activity (Fig. 14b), similar to the Early Jurassic palaeogeographical configuration. Research has indicated that Late Jurassic– Early Cretaceous depositional systems are characterized by braided stream –alluvial plain assemblages, in which strata of the Fenghuangtai Fm. mainly consist of paraconglomerates deposited by flood and debris-flow mechanisms along the southern margin of the Hefei basin. These depositional records probably indicate an increase in thrusting activity or topographic relief from late Middle Jurassic to early Late Jurassic–Early Cretaceous time, because no evident climate change can be found at this time. Later, the duration and intensity of uplift of the north Dabie Mountains even resulted in the erosion of Jurassic– Early Cretaceous strata at the southern margin of the Hefei basin with sedimentary depocentres migrating northward. This is why partial detritus of the upper section of the Zhougongshan Fm. in the northern Hefei basin lithologically is similar to that of the Jurassic– Early Cretaceous strata (Li, Z. et al. 1999). Early Cretaceous. Both sedimentary and structural results indicate basinwide changes in deposition between Late Jurassic–Early Cretaceous and Early Cretaceous time in the study area. Apparently, Early Cretaceous deposition in the southern margin of the Hefei basin mainly reflects the development of fan-delta –lacustrine systems. These fandeltas developed in response to extensional basin structures that were very different from those of Late Jurassic –Early Cretaceous time. Furthermore, Early Cretaceous deposition was mainly controlled by the Xiaotian –Mozitan normal fault with

Fig. 12. An explanatory north– south seismic profile across the Hefei basin, showing different structural frameworks between Jurassic–Early Cretaceous and Cretaceous–Tertiary strata (modified from unpublished data for the Shengli oilfield). The location of the seismic profile is shown in Figure 2b.

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Fig. 13. Isopach maps showing Mesozoic strata thickness in metres and depocentre changes in the Hefei basin system; compiled from Jia et al. (2001), Xu et al. (2002) and unpublished data for the Shengli oilfield. (a) Early Jurassic; (b) Middle Jurassic and Late Jurassic– Early Cretaceous; (c) Early Cretaceous; (d) Late Cretaceous.

northward dip rather than the Jinzhai –Huoshan – Shucheng thrust fault with southward dip (Figs 12 and 13c). In addition, a series of extensional faults, including southward-dipping faults in the north Hefei basin and westward-dipping faults in the eastern margin of the Hefei basin (Liu et al. 1993), restricted the depositional framework of this phase with disjunctive depocentres (Figs 13c and 14c). Late Cretaceous. Late Cretaceous depositional facies assemblages are characterized by sandy

braided river and alluvial plain facies showing decreased water depth. The Late Cretaceous depositional framework was also controlled by a set of extensional faults, particularly those with southward dip. As shown in Figure 13d, the Hefei basin chiefly consists of many half-graben style sub-basins mostly characterized by a northern bounding fault and depositional overlap towards the south (Fig. 12). Some larger-scale extensional margin faults (e.g. Xiaotian –Mozitan and Tan-Lu) were still active at this time (Lin et al. 2005).

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Fig. 14. Mesozoic sandstone content percentage isopach maps and regional tectonic palaeogeographical reconstruction of the southern Hefei basin system. (a) Early–Middle Jurassic; (b) Middle Jurassic and Late Jurassic– Early Cretaceous; (c) Early Cretaceous.

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Evidently dissimilar to other Late Mesozoic basins in North China, the Hefei basin evolution was principally is controlled by structures with approximately east–west strike. A set of structures with NE–SW and NNE–SSW strike had little effect on basin-fill of the Mesozoic, except the Tan-Lu fault along the eastern basin margin (Fig. 13).

Discussion Structural and volcanic evolution accompanying Mesozoic basin-fill Mesozoic basin-fill sequences of the northern margin of the North China Craton mostly start with volcaniclastic rocks, and pass upward through fine clastic rocks and mudstones of lacustrine facies and/or coal seams, to thick coarse clastic rocks and/or conglomerates in each basinfill phase, indicating a cyclic basin development (Figs 3 and 10). In the Early– Middle Jurassic to Middle–Late Jurassic cycles, the Yanshan –Liaoxi area principally experienced crustal thickening under a compressional tectonic regime, forming a mountain root with a thickness exceeding 60– 70 km (Deng et al. 1999), as shown by magmatic sources deeps than 50 km (Zhang et al. 2001). Consequently, coarse clastic alluvial – fluvial systems, such as the Tuchengzi Fm., developed under this compressional tectonic regime. In the Cretaceous tectonic cycle, however, the study area of the northern North China Craton was mainly characterized by an extensional tectonic regime and lithospheric thinning, although there is no evidence of typical bimodal igneous assemblages resulting from lithosphere rifting (Deng et al. 1999; Li, X. et al. 2004). In fact, published research has also suggested that the Mesozoic basins of the Yanshan –Liaoxi area were characterized by early magmatic diapirism and late compressional deformation and crustal thickening in each tectonic cycle (Deng et al. 1999, 2004). In addition, a structural evolution analysis shows that most basin-controlled structural strikes were jointly influenced by new structures of NE–SW and/or NNE–SSW strike as well as preexisting basement structures of east –west strike since the deposition of the Late Jurassic Tuchengzi Fm. (Li, Zi. et al. 2004). In other words, the inherited east –west structures were controlled by the Palaeoasian tectonic system. These structures were overprinted by the new NE–SW and/or NNE–SSW structures controlled by the circumPacific tectonic system of East Asia (Wang et al. 1990). Therefore we conclude that the Late Jurassic should be considered to be the tectonic transition phase in the Yanshan –Liaoxi area, on the northern margin of the North China Craton.

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Kinematic and chronological research on faults establishes the structural activity in the Yanshan– Liaoxi area (Zheng et al. 2000; Davis et al. 2001). Faults older than 180 Ma show southward thrusting; faults of 161–148 Ma (or 143–127 Ma) indicate northward thrusting or thrusting both northwards and southwards; and faults younger than 118 Ma or 116 Ma are principally attributed to an extensional event. A similar interpretation of tectonic evolutionary phases and attributes can be derived from sedimentary records, as described in this paper, although some minor geochronological differences still exist between structural and sedimentary stage limits. On the other hand, Early Jurassic volcanic assemblages containing basalts, and their lithogeochemical composition, show a wide geochemical range in the Yanshan–Liaoxi area, implying intensive crust –mantle interaction. However, intermediate –felsic and calc-alkaline volcanic assemblages of the Tiaojishan Fm., characterized by adakites with little basalt, do not show a wide change in neodymium and strontium isotopic composition and are characterized by low lead isotopic values and high potassium contents (Table 1), These characteristics are dissimilar to those of the Early Jurassic volcanic assemblages but similar to those of the crystalline basement and Mesozoic granites of the North China Block (Chen et al. 1995, 1997; Zhang 1995), indicating a close genetic relationship to the enriched mantle source of the North China continent. Similar to those of the Early Jurassic, the Early Cretaceous volcanic assemblages are characterized by widely varying geochemical compositions (Table 1), with basalts developed early. Therefore, we suggest that mantle sources partly influenced Early Jurassic volcanism. Mantle sources were relatively unimportant during Middle Jurassic volcanism, but significantly influenced Early Cretaceous volcanism in the study area. This shows that the tectonic transition to deep processes occurred in the Middle–Late Jurassic, earlier than the Late Jurassic transition phase marked by shallow basin structural evolution. These data probably indicate that an uncoupled relationship or time lag existed between the Mesozoic deep tectonic processes and the shallow responses in the Yanshan– Liaoxi area. Controlled by thrust faults with southward dips (Fig. 12), the Hefei basin was first filled by terrestrial clastic deposits of an alluvial system with multiple upward-coarsening units in the Jurassic–early Early Cretaceous. This depositional sequence reflects increasingly strong episodic thrust deformation in the northern Dabie Mountains and flexural depression to the north. From late Early Cretaceous time, twophase rifting processes are recorded by a set of lacustrine–delta systems (Fig. 10). Therefore, along the

43.1– 74.6 0.34– 2.72 12.4– 19.0 0.20– 5.40 3.40– 8.86 0.70525– 0.70690 0.511346– 0.521621 16.524– 18.252 15.252– 15.614 36.754– 38.730 Early asthenosphere – lithosphere interaction, late crust –mantle interaction

SiO2 (%) TiO2 (%) Al2O3 (%) K2O (%) Na2O þ K2O (%) 87 Sr/86Sr 143 Nd/144Nd 206 Pb/204Pb 207 Pb/204Pb 208 Pb/204Pb Geodynamic explanation

Geochemical composition is compiled from Li (1999) and Zhang (2005).

Early Jurassic Nandaling Fm. Xinglongdou Fm. Basalt – trachyte andesite – andesite

Lithostratigraphic age: Yanshan area: Liaoxi area: Volcanic assemblage: 49.0 – 81.0 0.08 – 1.75 11.5 – 20.0 0.67 – 5.92 4.21 – 10.9 0.70583– 0.70627 0.511700– 0.51184 16.370 – 17.37 15.19 – 15.38 36.19 – 37.01 Intensive crust – mantle interaction

Middle Jurassic Tiaojishan Fm. Tiaojishan Fm. Andesite – trachyte andesite

Table 1. Geochemical composition and geodynamic settings of the Mesozoic volcanism in the Yanshan –Liaoxi area

41.2 – 79.6 0.05 – 2.01 9.32 – 18.4 0.24 – 9.22 3.42 – 11.1 0.70474– 0.96470 0.51167– 0.512480 16.572 – 17.892 15.221 – 15.512 36.727 – 38.730 Early crust– mantle interaction, late asthenosphere– lithosphere interaction

Early Cretaceous Zhangjiakou Fm. Yixian Fm. Basalt – quartz andesite– rhyolite; trachyte– trachyte andesite

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southern margin of the North China Craton, basin-fill records show an evolutionary sequence from the Jurassic–early Early Cretaceous compressional regime to the late Early Cretaceous–Late Cretaceous extensional regime. Kinematics and chronology of faults reveal that the Xiaotian–Mozitan fault, separating the Dabie Mountains dome from the Beihuaiyang belt, mainly behaved like a normal fault with a northward dip that was active at 125–130 Ma (Faure et al. 1999; Lin et al. 2005), and controlled Cretaceous rather than Jurassic deposition as recorded in the Xiaotian sub-basin. However, an integrated interpretation of deep seismic profiles (Fig. 12) and depositional sequences show that the Jinzhai– Huoshan–Shucheng fault with its southward dip experienced early thrust activity that controlled the southern boundary of Jurassic deposition (Fig. 14a and 14b), and then was probably reactivated as an extensional fault influencing Cretaceous deposition. For the Hefei basin, two groups of volcanic rocks, including a 149 –138 Ma trachyandesite –trachydacite assemblage and a 132 –116 Ma basaltic trachyandesite –trachyandesite assemblage, existed in the late Jurassic –early Cretaceous basin-fill sequence (Fan et al. 2004). These volcanic assemblages are of distinctive geochemical composition and sources. Although this volcanism and its deep tectonic environments were generally considered as ‘continuous lithospheric thinning and decompressional melting’ by Fan et al. (2004), we infer that the older group of volcanic rocks was probably emplaced during a compressional tectonic regime that occurred at the same time as a phase of relaxation in the shallow crust. Our evidence for this interpretation is as follows: the older group of volcanic rocks contains much material from a crustal source different from the young group of volcanic rocks, and has a facies change with the Fenghuantai Fm. in which very thick alluvial conglomerates were deposited in response to active thrusting at the basin scale. In the northern foothills of the Dabie Mountains, basin-fills of the Cretaceous extensional regime are interstratified with 132– 116 Ma volcanic assemblages, indicating that extensional depositional systems in the shallow crust lagged behind the initiation of the ‘lithospheric thinning and decompressional melting’ suggested by Fan et al. (2004).

Basin geodynamic mechanisms around the North China Craton Basin-fill records around the eastern North China Craton reveal that Mesozoic continental basins in the Yanshan –Liaoxi and Hefei areas had similar evolutionary processes. The earliest depositional records of these basins, with similar structure

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style, are of early Jurassic age. These basin-fills principally record an evolutionary sequence from compression and lithospheric thickening before the Late Jurassic and/or Early Cretaceous to intracontinental stretching and lithospheric thinning after the Early Cretaceous, with similar tectonic transition phases in the late Jurassic. Along both southern and northern margins of the North China Craton, the late Mesozoic extensional structures occurred later in the shallow crust than the deep lithospheric thinning event. Lastly, similar alluvial deposits of great thickness are ubiquitous in the Mesozoic basins around the North China Craton. These coarse clastic sediments were deposited in response to Late Cretaceous extension. Nevertheless, there are also distinct differences in basin-fill records between the northern and southern margins of the North China Craton, as follows. (1) Volcanic rocks of varied compositions, including mafic, intermediate –mafic and intermediate –felsic assemblages, were developed through the basin-fill histories of the Yanshan–Liaoxi basins from Early Jurassic to Cretaceous time, but limited volcanic rocks of the calc-alkaline series filled the Hefei basin during the Late Jurassic– Early Cretaceous. These differences reflect distinct volcanic sources and deep tectonic processes between the northern and southern margins of the North China Craton. (2) Late Mesozoic lithospheric thinning began at about 163 Ma along the northern margins of the North China Craton, which finally resulted in intensive extension at basin scale with lithospheric thinning at 145–110 Ma (Zhai et al. 2004). Along the southern margins of the North China Craton, late Mesozoic lithospheric thinning seemed to begin at about 149 Ma, and intensive extension at basin scale occurred after 132 Ma. Therefore, compared with the northern margin of the North China Craton, the late Mesozoic tectonic transition occurred later along the southern margin. (3) Jurassic coarse clastic deposits in the studied basin-fill sequences suggest that fluvial systems with varied provenances developed along the northern margin of the North China Craton, but alluvial deposits of great thickness occurred along the southern margin of the North China Craton, implying more topographic relief along the southern margin than in the north. (4) Analysis of Mesozoic depocentre evolution suggests that a high degree of complexity accompanied frequent volcanism and high geothermal flow in the Yanshan– Liaoxi region, but depocentre migration from south to north accompanied by a relatively low geothermal flux took place in the Hefei basin. These differences in basin evolution probably resulted from different deep geodynamic driving mechanisms.

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A similar tectonic evolution from intracontinental compression to intracontinental extension apparently existed along both southern and northern margins of the eastern North China Craton in the Mesozoic. On the other hand, basin-fill records indicate distinctive basin geodynamic settings and possible processes: inhomogeneous crust –mantle interaction and intracontinental deformation along the northern margin of the North China Craton, compared with deep continental subduction of the Yangtze Craton and post-orogenic exhumation of the Dabie Mountains in the south. Therefore, we infer that the late Mesozoic tectonic transition was initially induced by crust –mantle interaction along the northern North China Craton, as recorded in Mesozoic basins in the Yanshan –Liaoxi area, and probably extended southward from there; that is, the late Mesozoic lithospheric thinning and subsequent tectonic transition should be a systematic process or tectonic cycle, which is independent of the Triassic plate convergence events around the North China Craton (Wang et al. 1990; Okay et al. 1993; Hendrix et al. 2001; Badarch et al. 2002).

Conclusions (1) An integrated comparative analysis of basin-fill, volcanic and structural records from the northern and southern margins of the eastern North China Craton indicates that the Mesozoic basins in the Yanshan –Liaoxi and Hefei areas had similar evolutionary histories. Basin-fill evolution occurred from compression and lithospheric thickening at Early–Late Jurassic time to intracontinental stretching and lithospheric thinning from Early Cretaceous time. A tectonic regime transition occurred during the late Jurassic –early Cretaceous, with a time lag between events in the shallow crust and deep lithosphere. (2) There were distinct basin evolutionary records between the two basin systems. First, many volcanic layers including mafic, intermediate – mafic and intermediate– felsic assemblages were developed in the Yanshan – Liaoxi basins from the Early Jurassic to Cretaceous. In contrast, limited calc-alkaline series filled the Hefei basin system during the Late Jurassic –Early Cretaceous. Second, late Mesozoic lithospheric thinning began at about 163 Ma and 149 Ma in the northern and southern margins, respectively, correspondingly resulting in structural extensional events at basin scale at about 145 Ma and 132 Ma. Third, coarse clastic deposition in the northern and southern basins during the tectonic transition phase is mainly characterized by fluvial and alluvial facies, respectively. These differences indicate greater topographic relief in the south than in the

north. Fourth, Mesozoic depocentre migrations are complicated in the north but organized into a trend from south to north in the Hefei basin system. (3) Basin-fill records of the southern margins of the eastern North China Craton were dominantly controlled by uplift and exhumation resulting from post-collisional orogenesis and post-orogenesis in the Dabie Mountains area. On the other hand, the northern basin evolution shows a control mechanism probably influenced by intensive crust –mantle interaction, with a tectonic stress transition from compression to stretching. (4) We infer that the late Mesozoic tectonic transition was first induced by crust –mantle interaction in the northern North China Craton and probably extended southward from there. Furthermore, Mesozoic lithospheric thinning and the subsequent tectonic transition was probably a systematic geodynamic process that had no relation to the Triassic plate convergence events around the North China Craton. This research was supported by the National Science Foundation of China (NSFC) (Grants 40234050, 40472069), Innovative Project of the Chinese Academy of Sciences (Grant KZCX1-07) and the National Key Fundamental Research Project (Grant G1999043303). Many sincere thanks go to F. Jin, R. Li, F. Yu and J. Zhang for their help in the field, and to M. Zhai and J. Qi for discussions regarding the structural framework of the study area. Finally, we wish to thank S.S Graham, S. Hendrix and B. Windley for their insightful and useful reviews of the scientific problems and revision to improve the English expression of the final version.

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LATE MESOZOIC TECTONIC TRANSITION X U , C., Q IU , L., L EI , M. & L I , X. 2002. The evolution of the Dabie orogenic belt based on the sedimentary styles and tectonic patterns of the Hefei basin, Anhui. Sedimentary Geology and Tethyan Geology, 22(2), 91–98 [in Chinese with English abstract]. X U , S., L IU , Y., J IANG , L., S U , W. & J I , S. 1994. Tectonic Regime and Evolution of Dabie Mountain. Science Press, Beijing, 13–19, 135–137 [in Chinese with English abstract]. Y ANG , Z. 1982. The characteristics of geological and tectonic evolution of the Tongbai– Dabie Mountain. Acta Geologica Sinica, (2), 123–135 [in Chinese]. Z HAI , M., M ENG , Q., L IU , J. ET AL . 2004. Geological features of Mesozoic tectonic regime inversion in Eastern North China and implication for geodynamics. Earth Science Frontiers, 11(3), 285–297 [in Chinese with English abstract]. Z HANG , L. 1995. Lithospheric Block Geology in the East Asia: Isotopic Geochemistry and Geodynamics of Upper Mantle, Basement and Granites. Science Press, Beijing [in Chinese]. Z HANG , Q., W ANG , Y., Q IAN , Q., Y ANG , J., W ANG , Y., Z HAO , T. & G UO , G. 2001. The characteristics and tectonic–metallogenic significances of the adakites

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Palaeogeothermal response and record of Late Mesozoic lithospheric thinning in the eastern North China Craton S. HU1, M. FU1, S. YANG2, Y. YUAN1 & J. WANG1 1

Institute of Geology and Geophysics, Chinese Academy of Sciences, Beijing 100029, China (e-mail: [email protected]) 2

China Offshore Oil Exploration and Development Research Centre, CNOOC, 100027, Beijing, China

Abstract: The palaeotemperature recorded by vitrinite reflectance (Ro) in the pre-Cenozoic uplifted stratigraphic strata, and in Palaeozoic–Mesozoic remnant basins outside the Cenozoic depocentres, has not been overprinted by later thermal events in the eastern North China Craton (NCC). Based on downhole Ro data from the Palaeozoic and the Mesozoic subsections, we reconstruct the temperature gradients when the subsections reached their maximum palaeotemperatures in the Middle Triassic and the Cretaceous, and calculate the corresponding heat flow histories since the early Mesozoic. The temperature gradient and heat flow were much higher in the Cretaceous (35–43 8C km21 and 73–83 mW m22, respectively) than in the Middle Triassic and at the present. The high palaeo-heat flow during the Late Mesozoic implies that the thickness of the ‘thermal’ lithosphere at that time was c. 65 km, about half the thickness of c. 135 km estimated for the Early Mesozoic. The change from a stable thermal regime to an active thermal regime took place during the Late Jurassic– Early Cretaceous (c. 110 Ma). This tectonothermal event was accompanied by extensive surface erosion, and is also evidenced in the areas adjacent to the NCC, such as the South Yellow Sea and East China Sea basins. Our study provides not only geothermal evidence for the Late Mesozoic lithospheric thinning, but also additional constraints on the thinning mechanism, which is currently being debated.

The North China Craton (NCC) is a major Archaean craton in China (Fig. 1). It was reactivated in the late Jurassic or early Cretaceous. The Late Mesozoic extension and widespread magmatism throughout the eastern part of China have been attributed to an eastward tectonic escape (Tapponnier et al. 1982), Pacific subduction (e.g. Yin & Nie 1996), lithospheric delamination and thinning resulting from the subduction (e.g. Gao et al. 1999, 2004; Wu et al. 2005), as well as the Siberia –Mongolia –Sino-Korean collision (e.g. Enkin et al. 1992). This substantial recent evidence (Menzies et al. 1993; Griffin et al. 1998; Menzies & Xu 1998; Gao et al. 1999, 2004; Fan et al. 2000) reveals that the thickness, the chemical composition, and the thermal state of the lithosphere in eastern North China experienced a dramatic transition during the Phanerozoic. Early Palaeozoic diamond- and xenolith-bearing kimberlites and Cenozoic mantle peridotite xenolithsbearing alkali basalts indicate that the early Phanerozoic lithosphere was thick and stable to depths within the diamond stability zone (180– 200 km) with refractory mantle mainly composed of harzburgite, whereas the Late Phanerozoic (post-Cretaceous) lithosphere was notably thinner (,80 km) with fertile mantle mainly composed of lherzolite (Song & Frey 1989; Menzies et al.

1993; Griffin et al. 1998; Menzies & Xu 1998; Xu et al. 1998; Fan et al. 2000; Xu 2001; Zhang & Sun 2002). The thermal state of this region in the Palaeozoic and Early Mesozoic was typical of a stable craton, in which the surface heat flow falls within a 40 – 50 mW m22 range (Xu 2001), whereas the Early Cenozoic lithosphere is characterized by a high heat flow (c. 80 mW m22) as inferred from the geothermobarometry of garnet-bearing peridotites and pyroxenites brought to the surface by alkali basalts (Xu, Y. et al. 1995; Xu, X. et al. 1998; Shi et al. 2000). In comparison with the present-day average heat flow of c. 63 mW m22 (Wang et al. 1985; Hu et al. 2000), the eastern NCC lithosphere must have been relatively warmer in the Early Cenozoic. The evidence mentioned above was mainly obtained from Palaeozoic and Early Cenozoic volcanic rocks as Mesozoic basalts are very scarce in the NCC. However, the Late Mesozoic is a key time period for understanding the large-scale tectonic movement, magmatic activity and change in tectonic regime in the eastern NCC (Deng et al. 1994; Zhao et al. 1994; Shao et al. 1997; Zhou et al. 2001, 2002, 2003; Zhang et al. 2002, 2003; Wilde et al. 2003; Zhai et al. 2004; Wu et al. 2005). Thus, any information on the Mesozoic thermal state of the lithosphere is of critical

From: ZHAI , M.-G., WINDLEY , B. F., KUSKY , T. M. & MENG , Q. R. (eds) Mesozoic Sub-Continental Lithospheric Thinning Under Eastern Asia. Geological Society, London, Special Publications, 280, 267–280. DOI: 10.1144/SP280.13 0305-8719/07/$15 # The Geological Society of London 2007.

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Fig. 1. Sketch map showing the geological setting of the boreholes used for pre-Cenozoic thermal history reconstruction in the eastern NCC.

importance to improving our understanding of the dynamic proceses of the Phanerozoic in the eastern NCC. The thinning of the lithosphere in the eastern NCC and its corresponding change of lithospheric

thermal structure should have affected the nearsurface temperature field in basins away from the magmatic zones, and been recorded by the thermal indicators in the strata of the Palaeozoic – Mesozoic remnant basins. The analysis of the

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palaeothermal regime of these basins can potentially provide direct geothermal evidence for the Late Mesozoic thinning of the lithosphere. Previous studies of the thermal history of the NCC have mainly been focused on the Cenozoic, as most of the samples were collected from boreholes in the Cenozoic basins (Zhou & Pan 1992), where the strata reached their maximum burial depth at the present and the earlier thermal history has been overprinted. Very few direct indicators of the Mesozoic thermal regime have been reported so far. In recent years, data from more deep boreholes penetrating to the Palaeozoic– Mesozoic strata have become available, as a result of the development of oil exploration. Many of those boreholes are located either in the Cenozoic uplift or in the Palaeozoic– Mesozoic remnant basin adjacent to the Cenozoic depocentres. We have found that distinguishable, fragmented palaeotemperature profiles for different subsections of a borehole are preserved (Hu et al. 2001; Fu et al. 2005), which means that the pre-Cenozoic strata experienced higher palaeotemperatures in the past. We have, therefore, estimated palaeotemperature gradients for each subsection in the deep boreholes by conversion of Ro to palaeotemperatures (Burham & Sweeney 1989). The two main questions we would like to address in this study are: (1) How did the palaeogeothermal regime respond to the supposed Late Mesozoic thinning of the lithosphere in the eastern NCC and adjacent areas? (2) With reference to the concept of ‘thermal’ lithospheric thickness, how was the thinning achieved?

Method A significant change or inversion in the thermal regime at shallow depths does not necessarily mean thinning of the lithosphere at greater depths, but a lithospheric thinning will certainly cause thermal inversion at shallow depths. The key problems are: is the record of the thermal inversion corresponding to the lithospheric thinning large enough to be detected, has it been preserved to the present? Some thermal indicators, such as organic matter, minerals and fluid inclusions, can record the palaeotemperature that the host formation experienced. However, not all the palaeotemperature records can be preserved to the present. Ro data are widely used in studies of basin palaeotemperature history as a maximum temperature recorder. This means that the host formation must experience no temperature higher than the previously recorded one for the fingerprint of the earlier thermal regime to survive. The preservation of such a maximum palaeotemperture record

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depends on a special combination of palaeotemperature gradient and burial depth. Continuous sedimentation or erosion is unlikely to lead to the recorded palaeotemperatures being preserved. Thick deposition and subsequent erosion can generally produce the required tectonothermal conditions for preservation of an earlier period of heating. In the eastern NCC, these tectonothermal conditions are likely to be met in the Cenozoic uplifted areas or in the Palaeozoic –Mesozoic remnant basins outside the Cenozoic depocentres. There are different approaches to thermal history reconstruction based on a kinetic model of thermal indicators, including stochastic inversion (Corrigan 1991; Lutz & Omar 1991; Gallagher 1995), the palaeo-heat flow method (Lerche et al. 1984; Hu et al. 2001) and palaeotemperature gradient-based methods (Duddy et al. 1991; Bray et al. 1992; Feinstein et al. 1996; Hu et al. 1998; O’Sullivan 1999). Taking into account the fact that the Paleozoic –Mesozoic basins experienced multiple episodes of tectonic movement and erosion as well as the complicated thermal history in the eastern NCC, we employ the palaeotemperature gradient method for thermal history reconstruction in this study. This method estimates the palaeogeothermal gradient from the maximum palaeotemperature profile determined in a vertical sequence of samples from a borehole. For a borehole section that had experienced multiple tectonothermal events, the whole sediment pile is divided into discrete structural layers separated by unconformities, but the structural layers can be combined into a united subsection in which all the strata experienced the maximum palaeotemperature at the same time (Fig. 2a). If different subsections experienced maximum palaeotemperatures at different geological times (i.e. an older (or lower) subsection was hotter than a younger (or upper) subsection (Fig. 2b), the palaeotemperature profiles for each subsection can be reconstructed independently based on the vitrinite kinetic models (Fig. 2c). A range of kinetic models are available for the conversion of vitrinite reflectance to palaeotemperature (Waples 1980; Wood 1988; Burham & Sweeney 1989). We apply the parallel chemical reaction model (Burham & Sweeney 1989) in the present study. Each Ro value can be integrated as a maximum palaeotemperature value if a heating path or heating duration is known or assumed. Estimates of maximum palaeotemperatures from Ro data were determined using the model proposed by Burham & Sweeney (1989, equation (2)). The palaeotemperature profile for each subsection can be constructed, and therefore, the palaeotemperature gradient (dT/dz) for each subsection can be obtained using a linear least-squares fit.

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Fig. 2. Method to determine the maximum palaeotemperature gradient and the eroded thickness. Assuming the borehole section (a) consists of three subsections (1, 2 and 3) these subsections reached their maximum palaeotemperature at different geological times (i.e. the present (t0), t1 and t2, respectively) during the burial process (b). Because an older (lower) subsection experienced a higher palaeotemperature, the palaeotemperature profiles of all subsections (c) were recorded and preserved. They provide distinguished palaeotemperature gradients (dT/dz)1, and the thickness (Ei) of the removed section on the corresponding unconformities is then obtained by dividing the difference between the surface temperature (Ts) and the intercept of the palaeotemperature profiles (Ti) at the top unconformity by the estimated palaeotemperature gradient.

The thickness of the removed section on the unconformities (Ei) can be estimated based on the estimated palaeotemperature gradients, the palaeotemperature at the unconformity (Ti) and the palaeo-surface temperature (Ts) (see Fig. 2): Ei ¼ (Ti 2 Ts)/(dT/dz)i. Using the reconstructed eroded thickness and the present-day thermal conductivity of the core samples, the burial and thermal maturation history can be reconstructed. The palaeo-porosity (e.g. Sclater & Christie 1980) and the palaeo-thermal conductivity for each formation can then be subsequently evaluated, and the palaeo-heat flow can be determined accordingly.

Procedure To be able to reconstruct the palaeothermal regime, a suitable site, where the palaeothermal record is likely to have been preserved, must be identified. The Palaeozoic– Mesozoic basins on the eastern NCC were fragmented in the Late Mesozoic and preserved either in the uplifted area as a remnant basin with thin Cenozoic sediments, or deeply buried beneath the thick Cenozoic sediments in the Cenozoic basins. Figure 3e shows the typical structure of the cratonic cover in the eastern NCC and the burial history for the Palaeozoic– Mesozoic strata. There are four basic styles of burial history (Fig. 3a–d). The Mesozoic palaeotemperature information is probably preserved, and can therefore be reconstructed, in the cases shown in Figure 3a and b. For the case shown in Figure 3c the palaeotemperature has been overprinted by a more recent maximum temperature as a result of continuous deposition. The case in Figure 3d is

between those in Figure 3a and c; the preservation of palaeothermal information in this case depends on whether the palaeotemperature gradient was high enough to have caused the Palaeozoic – Mesozoic strata to have experienced higher palaeotemperatures in the past than at present. Downhole data from 50 deep boreholes that penetrated to the Palaeozoic –Mesozoic strata in the eastern NCC were collected and analysed, but only 12 boreholes outside the Cenozoic depocentres have complete Ro profiles (Fu et al. 2005), especially for the Palaeozoic–Mesozoic strata, and have been selected for Mesozoic thermal history reconstruction (see Fig. 1). Six representative vitrinite profiles are shown in Figure 4, four of them from the eastern NCC, and two from the adjacent areas (South Yellow Sea and East China Sea basins; see Fig. 1 for their locations). The selected boreholes from the eastern NCC fall into two categories: (1) the Ro profile breaks into separated segments in different subsections, such as Yi135 and LD17-1-1z; (2) the Ro profile comprises only data from a single subsection, such as Deng 5 or Bogu 1. For the first category, it is important to note that the presence of a break in the vitrinite profile is significant in terms of the thermal history reconstruction, as the break is not caused by a fault but by an unconformity. The break means that the earlier (lower) subsection experienced higher palaeotemperatures than the later (upper) subsections, as the Ro data preserve a maximum palaeotemperature record and are irreversible. The break also indicates that the earlier palaeothermal regimes have been separately recorded by the segments of the Ro profile. The maximum palaeotemperature gradient

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Fig. 3. Schematic geological cross-section through the eastern NCC (e). Location is shown in Figure 1. There are four basic burial styles for the Palaeozoic–Mesozoic strata in the eastern NCC, and these are illustrated by the four burial cases (a–d).

can be inferred from the slope of the vitrinite profile segment in each subsection (see Fig. 2). The Ro value at the top of the subsection reflects the amount of section removed from the preserved subsection. The geological time when the subsection reached its maximum palaeotemperature should be somewhere within the time-gap of the unconformity marking the top of the subsection (i.e. the time when the erosion started). For example, the vitrinite profile for borehole Yi135 is separated into three segments (Li et al. 2001), corresponding to the Palaeogene, the Upper Jurassic–Lower Cretaceous, and the Upper Permian subsections (see Fig. 4), which means that these three subsections experienced their maximum palaeotemperatures at different geological times, and the lower subsections reached higher palaeotemperatures than did the upper subsections. The upper Permian subsection reached its maximum palaeotemperature during the Early Triassic to Middle Jurassic, the Upper Jurassic– Lower Cretaceous subsection during Late Cretaceous– Early Palaeocene, and the Palaeogene subsection during the Oligocene or at the present, as it may have been overprinted by a more recent thermal event. The offshore borehole LD17-1-1Z in the Bo Sea exhibits a similar pattern to Yi135; a broken Ro profile and higher Ro gradients in both Late Cretaceous and Palaeogene–Eocene subsections. For the second category, the Ro data might have recorded palaeothermal information or might represent only the present-day thermal regime, which

can be evaluated from the Ro values and the presentday burial depth or temperature. If the equivalent temperatures of the Ro values from Burham & Sweeney (1989) model are much higher than the present-day temperature at the corresponding depth, the Ro profile is a palaeothermal record, otherwise, the palaeothermal regime has been overprinted by the current thermal state and previous palaeothermal information has been erased from the Ro profile. In our selected boreholes, such as Deng 5, we only have Ro data from the Triassic to Lower –Middle Jurassic subsection. It is impossible to evaluate a palaeothermal history directly based on the relationship of the Ro profiles, because of the lack of Ro data in the overlain subsection. However, according to the measured Ro value at the top of Ro profile in this subsection, the Ro value of c. 0.75% at the present depth of 2.1 km, equivalent to a temperature of 135 8C from Ro –temperature conversion, is much higher than the present-day temperature (c. 69 8C) at this depth. The Ro profile, therefore, recorded the palaeothermal regime during the Middle Jurassic to Early Palaeogene. This illustrates that a single Ro segment can also provide useful palaeothermal information. Two boreholes located in the Cenozoic uplifted basins to the SE of the eastern NCC (CZ35-2-1 in the southern Yellow Sea basin and WZ26-1-1 in the East China Sea basin; see Fig. 1 for their locations) also show broken Ro profiles (Fig. 4), indicating a lower palaeotemperature gradient during the Early Triassic and a higher gradient

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Fig. 4. Six example borehole Ro profiles from the eastern NCC and adjacent areas (see Fig. 1 for locations). The straight-line segments are the subsections that experienced the maximum palaeotemperature at the same geological time. Dotted lines are unconformities.

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during the Late Mesozoic –Early Palaeogene (Yang et al. 2003). A similar pattern of changes in the thermal regime has been reported for the North Jiangsu basin (Yuan et al. 2005). This implies that the Late Mesozoic thermal inversion is not confined to the eastern NCC. The entire region of eastern China, including the offshore basins, may have been involved. We illustrate our thermal history reconstruction procedure in Figure 5, using the borehole of

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Yi135 as an example. From the borehole stratigraphic column (Fig. 5a), including lithology, age and unconformity, and the measured Ro profile, a preliminary evaluation can be made to assess whether the Ro data carry any valid palaeogeothermal information. If so, the Ro values are converted to maximum palaeotemperatures using the Burham & Sweeney (1989) model, and the palaeotemperature profile is constructed as shown in Figure 5b. The present-day temperature profile

Fig. 5. Procedure of thermal reconstruction illustrated by borehole Yi135, showing how Ro data can be used as a record of palaeothermal regime. The error bar represents an error of +10%. (See text for more detailed description).

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provides a comparison between the present and the palaeotemperature profiles in different subsections. The Cenozoic (Cz) temperature profile is consistent with the present-day temperature profile because of increased burial during the latest Tertiary. In contrast, the Mesozoic (Mz) and Palaeozoic (Pz) subsections have experienced much higher palaeotemperatures. The thermal regime of the Cz subsection has been overprinted by recent burial, whereas the maximum palaeotemperatures in the Mz and the Pz subsections have been preserved. The slope of the fitted linear relationship between the maximum palaeotemperature and the present depth provides a direct estimate of the maximum palaeogeothermal gradients for the Mz and the Pz subsections. A much higher palaeotemperature gradient was recorded in the Mz subsection than in the Pz and Cz subsections. The net amount of section removed on the unconformities can be obtained by dividing the difference between the palaeosurface temperature (assuming Ts ¼ 15 8C) and the intercept of the palaeotemperature profiles at the depths of the unconformities (Ti) by the estimated palaeotemperature gradient (see Fig. 2). Given the eroded thickness on the unconformities and assuming equal time for the deposition and erosion of the removed section, it is possible to reconstruct the burial history (Fig. 5c) and hence the palaeoporosity. Additionally, the palaeothermal conductivity during burial can be evaluated based on the presentday thermal conductivity measurements. This allows us to reconstruct the geotherms and palaeoheat flow histories at any given time, especially at the times when the subsections experienced their maximum palaeotemperatures as shown in Figure 5d, by assuming a linear variation of the estimated palaeotemperature gradients with time (Fig. 5e). As a final check, a comparison between the measured Ro values and the calculated Ro values, based on the T –t paths for the sampled depth and the forward model (Burham & Sweeney 1989), are performed to see if the reconstructed thermal history and the eroded thickness estimates are acceptable (Fig. 5f). All the above work relating to the thermal history reconstruction was completed using the program Thermodel for Windows (Hu et al. 1998, 2001).

Results Palaeotemperature gradient and palaeo-heat flow The results show that the temperature gradients since the Mesozoic have changed significantly in the eastern NCC. The gradient in the Early Mesozoic (c. 230 Ma) is typically ,30 8C km21 with a

mean of 26.9 + 3.3 8C km21; the Late Mesozoic was characterized by a much higher thermal gradient (35–43 8C km21 with a mean of 40.3 + 2.5 8C km21) than the present (31 –34 8C km21 with a mean of 32.4 + 1.1 8C km21) (Table 1). The temperature gradient during the Early Mesozoic was typical of a cratonic basin with relatively weak deep tectonothermal activity, whereas the gradient during the Mid –Late Mesozoic was as high as that in a typical active rifted basin with intense deep tectonothermal activity (Feinstein et al. 1996). Thermal conductivity in sedimentary cover generally increases with depth as a result of the compaction of sediments, which results in the temperature gradient decreasing with depth. The temperature gradient is, therefore, not an ideal parameter to represent the thermal regime of a basin. Heat flow is a more comprehensive parameter to characterize the basin thermal regime than temperature gradient, as heat flow is the product of thermal conductivity and temperature gradient. The palaeo-heat flow values based on the palaeotemperature gradient, measured present-day thermal conductivity and the reconstructed palaeo-porosity during burial for each borehole in the eastern NCC are listed in Table 1. Palaeo-heat flow values in the Late Mesozoic (73 –83 mW m22 with a mean of 78.9 + 2.9 mW m22) are much higher than that in the Early Mesozoic (47 –56 mW m22 with a mean of 52 + 3.5 mW m22) and that at the present (c. 63 mW m22) (Hu et al. 2000).

Denudation– erosion The thickness of the removed section on the unconformities is significant in the eastern NCC. The eroded thickness on the Late Jurassic– Early Cretaceous unconformity ranges from 0.9 to 3.9 km with a mean of 1.9 + 1.1 km. Another thick layer, of 1.5–5.2 km, was removed on the Mid– Late Triassic unconformity with a mean of 3.1 + 1.4 km. It should be pointed out that this erosion took place across the basins on the craton, not within the orogenic belt.

‘Thermal’ lithosphere The lower lithospheric boundary is an important structural and physical boundary within the interior of the Earth. It is geothermally defined as the boundary between the outer, conduction-dominated solid layer of the Earth, and the underlying convection-dominated asthenosphere (Morgan 1984). For a stable heat flow province, it is, therefore, possible to evaluate the lithospheric thickness from the steady-state conductive geotherm and the mantle solidus. A lithospheric thickness determined with such an approach is referred to as a ‘thermal’

Jiyuan depression Zhoukou Basin Mean

Nan 6 Zhoucan11

Xinzhu5 Guanshen1 Tanggu5 Deng5

Yi135 Yi155 Zhuang11 Bogu1 Changcan1 Sheng1

Jiyang depression

Changxian uplift Jizhong depression Linqing depression

Borehole

Region

34 31 32.4 + 1.1

57 56 60 + 2.6

58 63

62 58

33 31

32 33

59

Heat flow (mW m22)

32

Gradient (8C km21)

Present-day

35 40.3 + 2.3

41 40 41

40 41.3 41.3

43

Gradient (8C km21)

73 78.9 +

79 78 80

78 81 79

83

Heat flow (mW m22)

3.9 1.9 + 1.1

1.5 0.9 1.0

1.2 1.7 2.7

2.6

Erosion (km)

Middle – late Mesozoic (J3 – K1)

Table 1. Palaeotemperature gradient, heat flow and eroded thickness for boreholes in the eastern NCC

26.9 + 2.9

27

28

26

22.8 31

Gradient (8C km21)

52 + 3.0

56

55

47

52 52

Heat flow (mW m22)

Early Mesozoic (T2 – 3)

3.1 + 1.4

3.0

1.4

2.3

5.2 3.5

Erosion (km)

PALAEOGEOTHERMAL RESPONSE OF LITHOSPHERIC THINNING 275

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lithospheric thickness, to distinguish it from the lithospheric thickness determined by other geophysical methods (Pollack & Chapman 1977). This study adopted the dry basalt solidus as a reference mantle solidus to constrain the lower boundary of the lithosphere (Lachenbruch & Sass 1978): T ¼ 1050 þ 3z (where T is temperature in 8C, and z is depth in km). Ma et al. (1991) divided the crust of the NCC into three layers: the upper, middle and lower layers. The upper crust has been further divided into sedimentary cover and crustal basement. Together with the lithospheric mantle, a five-layer lithosphere model is adopted for analysis of the deep thermal structure in this study. In addition, the upper crustal sedimentary layer comprises Cenozoic, Mesozoic and Palaeozoic strata. We assume the thickness of the present-day crust to be 34 km, with the upper, middle and lower crust being 14, 9 and 11 km thick, respectively (Ma et al. 1991). The crustal thickness in the early Mesozoic is estimated to be 43 km, similar to that of the cratonic area without stretching, analogous to the Erdos Basin to the west. This means that the crust in the Late Mesozoic might have been significantly stretched; the lower crust, therefore, has been reduced to nearly half of the Early Mesozoic thickness, similar to that at present. The crust remains a five-layer structure but the thickness of the stretched lower crust and the sedimentary cover are adjusted accordingly (see Table 2). The thermophysical properties we used in this study are obtained from the weighted mean of the measured values for the sedimentary cover, and from references for the mid –lower crust and the lithospheric mantle (Pollack & Chapman 1977; Morgan 1984). Heat generation beneath the sedimentary cover is calculated using the exponential model: A(z) ¼ A0exp(2z/H), where A0 ¼ 1.26 mW m23 and H ¼ 8– 14 km depending on the crustal nature. The mean reconstructed palaeo-heat flow values (see Table 1) were used. The present-day heat-flow values are the means of the heat-flow measurements in this region (Hu et al. 2000). For the local Cenozoic rift basins on the eastern NCC, rifting continued until Oligocene time (c. 25 Ma) following the Late Mesozoic extension event and high palaeoheat flow persisted (e.g. Hu et al. 2001; Qiu et al. 2004). This Oligocene rifting does not appear to represent a regional event for the entire eastern NCC, but is local to the Cenozoic rifts. The calculated geotherms for the specific geological times and the corresponding mantle solidus are shown in Figure 6. The results indicate that the lithosphere in the eastern NCC experienced a significant thinning event in the Late Mesozoic, followed by a more

recent thickening. The lithospheric thickness of the NCC changed from 180–200 km in the early Palaeozoic (Griffin et al. 1998; Xu 2001) to c. 135 km in the early Mesozoic, to c. 65 km in the Late Mesozoic, and is now c. 75 km. The minimum thickness of c. 60 km is estimated for Oligocene (c. 25 Ma) for the local Cenozoic rift basins. The present-day lithospheric thickness inferred from the palaeothermal data is close to that estimated by other geophysical methods (Liu 1994). The ‘thermal’ lithosphere was thinned more than 50% from the Early Mesozoic to the Late Mesozoic. This geothermal evidence for the significant thinning of the eastern NCC corroborates that proposed by many other estimates based on geochemical work (e.g. Menzies et al. 1993; Griffin et al. 1998; Fan et al. 2000; Xu 2001). It should be pointed out that the uncertainty of the calculated lithosphere temperatures and ‘thermal’ lithospheric thickness could be large. First, the error in the reported surface heat-flow measurements is estimated to be about 10%, which can result in an error of about 50 8C in the calculated Moho temperature and 15 km in the lithospheric thickness, respectively. Second, the thermophysical properties, including heat production and thermal conductivity, were estimated based on empirical relationships and the errors are difficult to be quantify. Additionally, the results of this analysis may be affected by many other factors, such as the assumption of a steady-state lithospheric thermal regime during the Late Mesozoic, and the approximation of the upper mantle solidus by the dry basalt solidus. An uncertainty of 15% for the estimated lithospheric thickness can be expected (Pasquale et al. 1990). Nevertheless, the general trend should be reasonable.

Conclusions The eastern NCC experienced a thermal inversion (reactivation) the from low heat flow in the Early Mesozoic to high heat flow in the Late Mesozoic (locally to the Oligocene for the Cenozoic rift basins in the NCC). Such a significant change in the lithospheric thermal regime is consistent with the previously proposed lithospheric thinning in the Late Mesozoic. The thermal inversion is also observed in the North Jiangsu, South Yellow Sea and East China Sea basins adjacent to the NCC, suggesting that the Late Mesozoic tectonothermal event may have affected the entire region of Eastern China. The time of this inversion of the shallow thermal regime is estimated to be c. 110 Ma from the time gap (Late Jurassic–Early Cretaceous) of the unconformity corresponding to the regional transition of

Mz Pz Mz Pz Cz Mz Pz

3 2 5 2 5 2 2

Thickness (km) 2 3.0 2 3.0 1.6 2 3.0

K (W mK21)

Sedimentary cover

1.26 0.97 1.26 0.97 0.8 1.26 0.97

A (mW m23)

6

6

6

H (km)

2.3

2.3

2.3

K (W mK21)

Upper crust

9

10

11

H (km)

2.5

2.5

2.5

K (W mK21)

Middle crust

10

12

22

H (km)

2.5

2.5

2.5

K (W mK21)

Lower crust

K, Thermal conductivity; A, radiogenic heat generation; H, radiogenic thickness. Thermal conductivity and heat generation for the lithospheric mantle are 3.4 W mK21 and 0.03 mW m23, respectively.

Present

Mid – late Mz

Early Mz

Geological time

Table 2. Crustal structure and thermophysical properties for specific geological times in the eastern NCC

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and the significant erosion at the surface as recorded by the thermal indicators. This research was supported by the Chinese Academy of Sciences (Grant KZCX1-07), and the National Science Foundation of China (Grant 40172100). This paper was greatly improved by the careful review and useful comments of R. Brown. We would like to thank X. H. Zhou and S. P. Huang for their valuable input into the ideas presented in this paper. We are grateful to SINOPEC, PetroChina and CNOOC for providing downhole data used in this paper. The data used in this paper are available from the author upon request.

Fig. 6. Estimated geotherms in the lithosphere for different geological times and their relationship with the mantle solidus, dry basalt solidus (BDS). The base of the thermal lithosphere is defined as the intersection of the model geotherms with the BDS. The error for the estimate of lithospheric thickness is assumed as +15%.

the thermal regime. The transition in the shallow thermal regime was probably preceded by a deep tectonothermal event. Our palaeotemperature analysis shows that the lithosphere in the eastern NCC changed its thickness dramatically from c. 135 km in the early Mesozoic to c. 65 km in the Late Mesozoic (c. 110 Ma). In addition, the Late Mesozoic thermal reactivation was accompanied by extensive surface erosion, providing tectonothermal constraints on the mechanism of lithospheric thinning in the eastern NCC. The thermal impact from the Late Mesozoic lithosphere thinning has at least been partially relaxed since the Cenozoic. The present-day lithospheric thickness of the eastern NCC is c. 75 km. Different NCC lithosphere thinning mechanisms have been proposed by several research groups. Some researchers have proposed that the thinning could have resulted from the replacement of the stable cratonic lithospheric mantle during the Paleozoic oceanic mantle in the Mesozoic –Cenozoic (Zheng et al. 2001), or from a major transformation of subcontinental lithosphere (Zhou et al. 2001, 2002, 2003; Zhang et al. 2002, 2003; Zhou & Zhang 2006). Some others have attributed the thinning to the replacement of the lower crust in the Mesozoic– Cenozoic (Zhai & Fan 2002), or mantle delamination (Gao et al. 1999, 2004; Wu et al. 2005). Deng et al. (1994, 2004) suggested unrooting of the orogenic belt as the major cause for the thinning. The present study offers an important new constraint to the continuing debate concerning the NCC lithosphere thinning mechanism: a successful candidate must be able to reasonably explain both the inversion of the thermal regime

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East Marginal Fault of the Yellow Sea: a part of the conjunction zone between Sino-Korea and Yangtze Blocks? T.-Y. HAO1, Y. XU1, M. SUH2, Y. XU1, J.-H. LIU1, L.-L. ZHANG1 & M.-G. DAI1 1

Institute of Geology and Geophysics, Chinese Academy of Sciences, Beijing, PO Box 9825, 100029, China (e-mail: [email protected]) 2

Kongju National University, Kongju 314-701, S. Korea

Abstract: The conclusions of most previous studies on the eastward extension of the collision zone between the Sino-Korean and Yangtze Blocks can generally be divided into two categories: (1) the collision zone is connected to the Imjingang belt and crosses the Korean Peninsula; (2) the eastward extension of the collision zone does not enter the Korean Peninsula, nor is connected to the Imjingang belt. Recent geophysical studies on gravity and velocity tomography in the Yellow Sea and adjacent regions, respectively, have provided geophysical evidence for a nearly north– south-trading dextral strike-slip fault in the eastern margin of the Yellow Sea, which we have named the East Marginal Fault of the Yellow Sea (EMFYS). The geophysical evidence indicates that the EMFYS extends to great depth. It dips westward, and within 100 km depth the dip angle is very steep. The geophysical characteristics on its two sides show that they belong to different tectonic units. This fault is connected to the Wulian– Qingdao Fault in the north, and to the South Marginal Fault of Jeju Island (SMFJI) in the south (the SMFJI is a part of the collision zone). Therefore the EMFYS is considered as a part of the junction zone between the Sino-Korea and Yangtze Blocks, with the Korean Peninsula being part of the Sino-Korea Block. In the Late Triassic, it is inferred that dextral strike-slip movement took place on the EMFYS, and in the same geological period sinistral strike-slip took place on the Tan-Lu Fault Zone. Under north– south tectonic stress the Yangtze Block was translated northward and inserted into the Sino-Korea Block. Therefore the junction zone between the two blocks formed a gigantic Z-shaped tectonic belt. According to the gravity data and seismic tomography it is also inferred in this paper that the junction zone between the Yangtze and Cathaysia Blocks extends from the Jiangshao Fault eastward to the southern edge of the Hida Block, Japan.

The Yellow Sea between the Chinese continent and the Korean Peninsula spans two tectonic blocks, the Sino-Korea Block (or North China Block) and the Yangtze Block. It is an important place for understanding the evolution process of the Asian continental margin. As the collision zone between the Sino-Korea and Yangtze Blocks, the Qinling –Dabie collision zone has been extensively studied, and many geological and geophysical results have been published (Yuan et al. 1994, 1997; Wang et al. 2000; Zhang et al. 2000). This zone is cut by the Tan-Lu Fault (TLF) then connected to the Zhucheng–Rongcheng collision zone (Wang et al. 2000; Wan 2004). However, there is much controversy concerning the whereabouts of this collision zone eastward of the Yellow Sea. Yin & Nie (1993) suggested that this zone extends eastward across the Korean Peninsula to connect with the Honam shear zone. Other researchers have proposed that this collision zone is connected with the Imjingang orogen to form the suture between the Sino-Korea and Yangtze Blocks (Yang 1989; Li et al. 1992; Cai 2002; Cai et al. 2002).

However, in recent years, with more intensive studies of Sino-Korean comparison, the location problem of the Sino-Korea and Yangtze collision has again been given attention by geoscientists. Oh et al. (2005) suggested that the southwestern part of the Gyeonggi massif is an extension of the Triassic Dabie –Sulu collision belt in China. Oh et al. and Seo et al. suggested that the Dabie – Sulu collision belt in China extends to the Hongseong–Odesan belt in Korea (Oh 2006; Oh & Kusky 2006; Seo et al. 2005). Recently, some geoscientists have searched for geophysical evidence of the collision belt in the Yellow Sea region. Choi et al. (2006) suggested that the suture zone continues in an ENE direction to the region of the Imjingang Belt and Hongseong in the western Korean Peninsula, based on interpretation of the GRACE satellite gravity dataset. On the basis of gravity and seismic tomography, Hao et al. (2002) proposed that there is a dextral strike-slip fault in the eastern margin of the Yellow Sea. In that study, the gravity data used were taken from the Bouguer gravity database (Liu 1992), and the 5 km  5 km grid data were obtained through

From: ZHAI , M.-G., WINDLEY , B. F., KUSKY , T. M. & MENG , Q. R. (eds) Mesozoic Sub-Continental Lithospheric Thinning Under Eastern Asia. Geological Society, London, Special Publications, 280, 281–292. DOI: 10.1144/SP280.14 0305-8719/07/$15 # The Geological Society of London 2007.

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Mercator projection and interpolation. The wavelet transform of gravity data was also determined in that study. Hao et al. believed that the dextral strikeslip fault in the eastern margin of Yellow Sea is connected with the Wulian–Qingdao–Rongcheng collision zone in the north, and extends southward to the southern margin of Jeju Island, where it joins the south. In our Marginal Fault of Jeju Island (SMFJI) earlier papers (Hao et al. 2002, 2003a, b) this fault was called the West Margin Fault of the Korean Peninsula. However, in a later study on deep structures of the Yellow Sea and East China Sea (Hao et al. 2006), we renamed it the East Marginal Fault of the Yellow Sea (EMFYS), which tallies better with the actual situation. The knowledge of the EMFYS supports the tectonic framework in which the whole Korean Peninsula belongs to the Sino-Korea Block (Lee et al. 1998; Chang 2000; Wan 2001; Ishiwatari & Tsujimori 2003), and the EMFYS is considered as part of the junction zone between the Sino-Korea and Yangtze Blocks. In the present paper we summarize different opinions on the eastern extension of the Sino-Korea and Yangtze collision zone, and provide geophysical evidence for the location of the junction zone in the Yellow Sea.

Knowledge of the eastern extension of the Sino-Korea and Yangtze collision zone Zhang (1986) pointed out that the collision zone between the Sino-Korea and Yangtze Blocks extends from the Jiashan–Xiangshui Fault eastward. After entering the Yellow Sea it extends first to the NE, then turns southward with a NNW–SSW trend to the southern end of the Korean Peninsula. The NNW–SSE structure is reflected as the boundary of magnetic anomalies (Fig. 1). Zhang (1986) also thought that the Korean Peninsula is a relatively uplifted tectonic unit, because ‘in the north and south of Korean Peninsula, the Palaeozoic cover is identical with North China’. Zhang (1986) considered that the major part of the Korean Peninsula belongs to the Sino-Korea Block. Although they acknowledged that ‘Silurian fossils and Devonian coral fossils were found in the gravel of Jurassic conglomerate rocks in Songnim and other places’, they thought these were probably products of the Imjingang depressions. Liu (1992) put forward a new model for the eastward extension of the Sino-Korea and Yangtze junction zone in the Geological and Geophysical Series Maps of China Seas and Adjacent Regions. He thought that to the east of the TLF the Wulian–Qingdao Fault extends eastward and connects with the Imjingang Fault Belt (IFB),

forming the collision zone of the two major blocks (i.e. Qinling– Dabie– Qingdao–Wulian – Imjingang). Li et al. (1992) pointed out that the evidence of the location of this collision zone came mainly from the discovery of the Imjingang group. The strata of the Imjingang group are as thick as 2100– 3600 m, and contain a large quantity of coral, cephalopod and crinoid fossils, which are Devonian palaeontological groups of South China type. However, Li et al. (1992) also admitted that the lithology and lithofacies of the Imjingang group are different from those of strata in the corresponding position in the Yangtze Block. Yin & Nie (1993) proposed that the collision zone extended eastward across the Korean Peninsula and connected with the Honam shear zone. This tectonic model is widely accepted. They not only made a comparison in terms of sedimentation and lithofacies, but also gave a tectonic model for the processes related to the collision of the North and South China Blocks in the Late Carboniferous to Early Jurassic period. In particular, they also considered the evolution and effect of the TLF and made a comprehensive analysis of various influencing factors. The results of later studies on the eastward extension of the collision zone can be generally divided into two categories: (1) the eastward extension of the collision zone is connected to the IFB and crosses the Korean Peninsula; (2) the eastward extension of the collision zone does not enter the Korean Peninsula, nor is it connected with the IFB. These studies respectively provided geological and geophysical evidence to support their conclusions (Qin et al. 1989; Cai 2002; Feng et al. 2002). In recent years many researchers have attempted to compare China and Korea in studies of the Sulu ultrahigh-pressure (UHP) metamorphic belt, or have used geophysical information on structures to search for evidence of the eastward extension of the junction zone. Lee et al. (1998) indicated that, in view of the overall similarity of the age, geology, lithology and geochemistry, the whole Korean Peninsula should belong to the Sino-Korea Craton. Some geochemical results (Teraoka et al. 1999) also indicated that the Sulu suture belt does not extend to Korea, but turns southward beneath the Yellow Sea. Oh et al. (2005) found an eclogitefacies metamorphic event in South Korea and believed that it is an extension of the Triassic Dabie –Sulu collision belt in China. However, observations by Zhai & Guo (2005) indicated that the eclogites have foliation with a north–south trend. Wan (2001) pointed out that the Gyeonggi Block, south of the Imjingang River, ‘is possessed of the typical crystalline basement, Lower Palaeozoic sedimentary rock formations, and the

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Fig. 1. Tectonic subdivisions and sutures between the Sino-Korea and Yangtze Blocks, after Zhang et al. (1986).

disconformity between Middle Ordovician and Carboniferous systems of Sino-Korea Block’. Therefore he inferred that ‘the main part of southern Korean Peninsula is still a part of Sino-Korea Block’. Chwae et al. (2005) expressed the same opinion. He pointed out that UHP minerals similar to those of the Dabie –Sulu metamorphic rocks of Eastern China were not found in the Imjingang or Ogcheon orogen, and the Imjingang orogen did not cross the Korean Peninsula from east to west; instead it bent northwards (Chwae et al. 2005).

Geophysical evidence for the East Marginal Fault of the Yellow Sea Geophysical surveying by chinese researchers in the Yellow Sea area began in the 1960s (Liu 1992). More detailed studies were the basic surveys for the Geological and Geophysical Series Maps of China Seas and Adjacent Regions (Liu 1992). The conclusion of that research project was that the collision zone of the Sino-Korea and Yangtze Blocks extends eastward and connects with the IFB, then crosses the Korean Peninsula to the east (Liu 1992). It was also concluded that another collision zone in the Yellow Sea, between the Yangtze and Cathaysia Blocks, starts as the Jiangshao Fault in East China, runs northeastward and connects with the Kwangju Fault in the

Korean Peninsula, which also crosses the peninsula (Fig. 2). According to this interpretation the Korean Peninsula is divided into three parts, belonging to the Sino-Korea, Yangtze and Cathaysia tectonic units. In the vicinity of Qingdao–Rizhao, the Wulian– Qingdao– Imjingang orogen is manifested mainly as linear magnetic anomalies, which are caused by granite masses intruded along the large-scale fault belt (Huang & Wang 1992). In 1992 we studied the Bouguer gravity anomalies on a 5 km  5 km grid. By analytical continuation, wavelet transformation and other processing techniques we extracted linear structural information on the Yellow Sea area, and defined the location of a nearly north–south-trending dextral strike-slip fault in the eastern margin of the Yellow Sea (Hao et al. 2002). This fault belt is named the EMFYS. It is a boundary between regions of different geophysical characteristics. In particular, on the two sides of the fault the nature of the gravity field and Moho depth distribution are very different (Fig. 3). The depths of the Moho discontinuity were calculated using the harmonic series method (Jiang 1989), taking 30 km as the standard crustal thickness and 20.43 g cm23 as the density difference. The EMFYS cuts the normal trends of the Moho isobaths: on the western side the dominant extension direction of the isobaths is NE–SW or ENE –WSW, whereas on the eastern side they are basically along the

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Fig. 2. Simplified sutures of East China and the adjacent seas after Liu (1992). Continuous line indicates a fault; dashed lines indicate block junction belts. The Korean Peninsula is divided into three parts: the Sino-Korea, Yangtze and Cathaysia Blocks, from north to south.

shoreline of the Korean Peninsula, displaying different characteristics (Hao et al. 2003b). The detailed shadowgraph of the first-order gravity wavelet shows that this fault has the dextral strikeslip character. More importantly, the result from gravity studies indicates that the Wulian– Qingdao Fault (WQF) does not enter the Korean Peninsula. Like most other NE–SW-trending linear anomalies in the western Yellow Sea, it is also cut off to the west side of the EMFYS (Fig. 4). In addition, the P-wave velocity profile along 388N (Hao et al. 2002) reveals velocity structure differences on the two sides of the EMFYS, and confirms that this fault penetrates to a great depth. Because of the limited resolution in depth of the velocity inversion, the character of the EMFYS in the upper mantle cannot be well defined. A recent research result concerning the deep structure of the marginal seas in Eastern China (Xu et al. 2006) has made up this deficiency: on the P-wave velocity image for 77 km depth

(Fig. 5), the EMFYS is clearly shown. Along the western margin of the Korean Peninsula a narrow low-velocity belt corresponds to the location of the EMFYS. Its strike direction and length are consistent with the result from gravity data (Hao et al. 2002). The velocity profile crossing the southern end of the Korean Peninsula (Fig. 6; Xu et al. 2006) shows that above 100 km the fault is very steep, and below 100 km it dips towards the west. It should be noticed that the P-velocity image of the Korean Peninsula on the east side of the EMFYS shows high velocity above 200 km, which is obviously different from the velocity distribution in the southern Yellow Sea on the west side of the EMFYS. However, high velocity characterizes shallow depths, but low velocity is characteristic of depths below 100 km. Therefore the two sides should belong to different tectonic units. The SMFJI (collision zone) that is connected to the EMFYS in the vicinity of Jeju Island is also a deep penetrating tectonic belt (Hao et al. 2003b).

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Fig. 3. Distribution of Moho discontinuity depth in the Yellow Sea and adjacent regions. EMFYS, is East Marginal Fault of the Yellow Sea; SMFJI, South Marginal Fault Zone of Jeju Island. The Moho discontinuity depths were calculated using the harmonic series method (Jiang 1989), taking 30 km as the standard crustal thickness and 20.43 gcm23 as the density difference.

In the southern margin of Jeju Island it extends in a NE–SW to ENE –WSW direction towards the Tsushima Straits (Figs 4 and 5). It dips towards the north and extends to about 230 km depth, indicating that it is a lithospheric-scale fault (Fig. 6). Park et al. (2005) mentioned that the mantle under Jeju Island has similar geochemical characteristics to the mantle under the South China Block. This is consistent with the result of Hao et al. (2002). Because the SMFJI is a north–south trending collision belt, the mantle under Jeju Island should

belong to the Yangtze Block (in Park’s paper, the South China Block is named the Yangtze þ Cathaysia Block).

A conjecture on the collision zone between the Sino-Korea and Yangtze Blocks In view of the above geophysical results, there is no evidence indicating that the Sulu orogen or WQF in Eastern China entered the Korean Peninsula; on the

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Fig. 4. Shadowgraph of first-order gravity wavelet details produced by discrete wavelet transform (from Hao et al. 2002, reproduced with permission of Chinese Journal of Geophysics).

contrary, they turned to the south of the Yellow Sea. Gravity data indicate that the characteristics of the gravity field on the two sides of the EMFYS are very different. The gravity field characteristics of the Korean Peninsula are different from those of the southern Yellow Sea, but are similar to those of NE China. The Moho depth

distribution is also divided by the EMFYS; the trends of isolines to either side of the EMFYS are very different, implying that they belong to different tectonic units. P-wave velocity images show a clear trace of the EMFYS above 100 km, and the velocity structures on the two sides of the EMFYS are different.

Fig. 5. Velocity perturbation image at 77 km depth and inferred junction zone between the Sino-Korea and Yangtze Blocks, and the location of the profile shown in Figure 6. The velocity perturbation image is from Xu et al. (2006, reproduced with permission of Chinese Journal of Geophysics).

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Fig. 6. Velocity profile L –L0 crossing the southern end of the Korea Peninsula, after Xu et al. (2006, reproduced with permission of Chinese Journal of Geophysics). The profile location is shown in Figure 5.

Summarizing the above facts, we consider that the EMFYS is a part of the junction zone between the Sino-Korea and Yangtze Blocks. As a result of north–south-trending compression, the TLF in Eastern China underwent sinistral strike-slip in the Late Triassic (Yin & Nie 1993; Wan et al. 1996; Wang et al. 2000). It is conjectured that in the same geological period dextral strike-slip took place on the EMFYS, and the Yangtze Block was translated northward into the Sino-Korea Block, forming a gigantic Z-shaped indented structure (Fig. 7). Therefore, the Korean Peninsula should be a part of the Sino-Korea Block.

Geophysical evidence and inference respecting the junction zone between the Yangtze and Cathaysia Blocks According to the above statement, if the southern segment of the Sino-Korea and Yangtze junction zone runs eastward along the Jeju collision zone, then the location of the junction between the Yangtze and Cathaysia Blocks should also be reconsidered. Based on geomagnetic data for the Yellow Sea, Wang et al. (2003) inferred that there is a continuous NE– SW striking deep structure in the southern Yellow Sea, which he ascribed to the Fujian–Lingnan uplift. This inference is basically consistent with the tectonic division scheme of Zhang et al. (1986). Based on the Bouguer gravity data and P-wave tomographic result, Hao et al. (2003a) deduced that this junction zone should extend from Hangzhou Bay through the Tsushima Straits to north Japan. The supporting evidence is as follows: (1) the characteristics of gravity fields on the two sides of this tectonic belt are different;

(2) the tectonic belt itself is a clear linear anomaly; (3) it is a boundary between regions of high and low velocities (Figs 5 and 6). Structural studies support such a division. Tsujimori et al. (2000) found eclogitic blueschist in the SE edge of the Hida Block, Japan (i.e. in the outer zone of the Hida Block). Wan (2004) deduced that the location of blueschist should be a collision zone, and considered that the Hida Block belongs to the Sino-Korea Block. Whereas the outer zone of the Hida Block corresponds to the Dabie collision zone, the inner zone of the Hida Block is comparable with the Yangtze Block. Therefore, the junction zone extends to the east in the vicinity of Jeju Island, and passes the southern edge of the Hida Block. The Yangtze Block may become very narrow in the east (Fig. 7). However, doubt still exists for us. Why, along the boundary belt between the Yangtze and Cathaysia Blocks, can we not find the Palaeozoic subduction complex as occurs in Japan? Why do we not find evidence of collision along the SMFJI? According to the tectonic evolution history of China, at the end of the Early Palaeozoic era, the Qinling area underwent a Palaeozoic subduction at about 400 Ma (Wan 2004). This subduction mostly occurred in the eastern Qinlin area (the present-day Shanxi and western Henan provinces in China). There is evidence of subduction in the Hida area similar to that in the Qinling area, but no evidence of subduction has been found in the Qinling–Dabie collision zone, especially in Shandong (i.e. the Sulu area). Therefore, we infer that a collision complex might not exist there. Moreover, it is found that the ophiolite zone of c. 1000 Ma is reduced by subduction in some areas of the Qinling–Dabie belt. Here, strong deformation

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Fig. 7. Distribution of the junction zone between the blocks we propose in East China and the adjacent sea areas. Bold lines indicate the junction zones, and arrowheads show the direction of movement of the blocks. Fine lines are fault zones. The dashed line is a conjectural fault. TLF, Tancheng– Lujiang Fault Zone; WQF, Wulian–Qingdao Fault; EMFYS, East Marginal Fault of the Yellow Sea; SMFJI, South Marginal Fault of Jeju Island. It is inferred that the gigantic Z-shaped indented structure formed in the late Triassic and was composed of the WQF, EMFYS and SMFJI.

occurred mainly on the northern side of the Qinling mountains (Wan 2004). Magmatism and medium– deep metamorphism around 1000 Ma can also be found in the Dabie mountains area. However, ophiolites and even evidence of bimodal volcanism exist only in some parts of the Qinling area but not in the Jiaodong area. Therefore, we conclude that it is difficult to find evidence of collision in both Hida and Shandong, and that further surveying and research is needed in both areas.

Summary The geophysical evidence indicates that the EMFYS along the eastern margin of the Yellow Sea is a dextral strike-slip fault extending to great depth. It dips westward, and within 100 km depth

the dip angle is very steep. The geophysical characteristics on its two sides show that they belong to different tectonic units. This fault is connected to the WQF in the north, and to the Jeju collision zone in the south. Therefore it is considered as a part of the collision zone between the Sino-Korea and Yangtze Blocks, and the Korean Peninsula should belong to the Sino-Korea Block. In view of the fact that the TLF underwent sinistral strike-slip faulting in the Late Triassic, it is inferred that dextral strike-slip took place on the EMFYS in the same period. Under north–south tectonic stress the Yangtze Block was translated northward and inserted into the Sino-Korea Block. Therefore the collision zone between the two blocks formed a gigantic Z-shaped tectonic belt. According to the gravity data and seismic tomographic results it is inferred that the collision zone between the Yangtze and Cathaysia Blocks

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extends from the Jiangshao Fault eastward to the southern edge of the Hida Block of Japan. If this tectonic model is correct, then the Yangtze Block may become very narrow to the east. We are grateful to G. D. Liu and T. F. Wan for their help with tectonics. We also thank K. H. Chang for his constructive discussion. Geophysical research on the Yellow Sea area is supported by NSFC project 40074021, 40674046, 40620140435 and the National Key Basic Research Development Program’s project G20000467-01.

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Crustal P-wave velocity distributions and metallotectonics around the North China Craton X. CHANG1, Y. LIU1, M.-G. ZHAI1 & Y. WANG2 1

Institute of Geology and Geophysics, Chinese Academy of Sciences, Post Box 9825, Beijing 100029, and College of Geophysics and Oil Resource, Yangtze University, China (e-mail: [email protected]) 2

Department of Earth Science and Engineering, Imperial College, South Kensington, London SW7 2BP, UK Abstract: We have collected seismic data, performed high-resolution seismic tomography in the North China Craton and analysed the relationships between crustal seismic velocity distributions and regional tectonics and metallogenesis. In the upper and middle crust velocity anomalies are distributed along east–west- and NNE–SSW-trending structures. Most belts of Cenozoic mineral deposits including gold in the North China Craton coincide with high-velocity anomalies, and the North China basin coincides with a low-velocity zone. Compared with the upper crust, the low-velocity anomalies in the lower crust are diffuse and extensive, which suggests that hightemperature material has upwelled from the mantle. High-temperature material in the lower crust provided buoyant, hot, mineralizing fluids that uplifted and formed the Cretaceous mineral deposits in the upper crust.

Seismic velocity depends on pressure, hydrostatic temperature and composition (Gardner & Gregory 1974; Gregory 1976), and thus can provide direct evidence for interpretion of the structure of the crust or mantle (Benz et al. 1996; Hauksson & Haase 1997; Ghose et al. 1998). To map velocity distributions, seismic tomography has been extensively employed (Thurber & Aki 1987; Humphreys & Clayton 1988; Inoue et al. 1990; Spakman et al. 1993; Liu et al. 2005). For example, in the North China Craton, seismic velocity tomography has been used as an aid to understand the tectonic evolution, earthquake mechanisms, orogenesis and active faults (Shedlock & Roecker 1987; Liu et al. 1990; Zhu & Zeng 1990; Sun & Liu 1995; Liu & Chang 2001; Wang et al. 2003; Huang & Zhao 2004). In contrast, mineralogists have paid studied the mineral resource belts in the North China Craton by deducing ore-forming mechanisms based on structure and mineral samples gathered on the surface (Zhai 1998; Zhai & Yang 2001; Zhai, M. et al. 2002; Nie et al. 2004; Zhai, Y. et al. 2004). Although seismic tomography is able to provide information on subsurface structure, few seismic tomographic studies have been related to ore-forming mechanisms. Physical properties such as density and temperature of minerals and rocks can be related to seismic velocity variations. In this paper, we report new data on high-resolution seismic tomography in the North China Craton, velocity images and structures, and

interpret the relationship between seismic velocity distribution in the crust and metallotectonics. The eastern North China Craton is demarcated by three faults (Fig. 1): the north boundary fault, the south boundary fault, and the Tanlu fault in the east (Cui & Li 1990; Cheng 1994; Wang et al. 2005). The North China Craton underwent orogenic movements in the Proterozoic and strong tectonic– igneous activity in the Mesozoic (Ye et al. 1985, 1987; Kern et al. 1996; Menzies & Xu 1998; Fan et al. 2000). Major faults and tectonic –igneous and metamorphic belts crop out along the eastern NNE –SSW-trending boundary and the two east – west-trending boundaries of the craton (Cheng 1994; Wang et al. 2005), and they contain various major mineralization zones (Zhai & Yang 2001; Zhai et al. 2002) including major gold deposits, as indicated in Figure 1b (Chen et al. 1998; Zhai 1998; Deng et al. 2003; Nie et al. 2004; Zhai et al. 2004). Rb–Sr isochron ages for pyrite (Yang & Zhou 2001), sensitive high-resolution ion microprobe (SHRIMP) U –Pb ages of zircons from wall rocks and dykes that crosscut the ore bodies (Wang et al. 2001), and 40Ar/39Ar ages of alteration minerals such as sericite (Zhang et al. 2003) indicate that the gold mineralization was mainly generated in the Late Mesozoic (Yang et al. 2003). Accordingly, we use seismic velocity distributions to clarify the tectonic framework that relates to the evolution and distribution of mineral deposits in the region.

From: ZHAI , M.-G., WINDLEY , B. F., KUSKY , T. M. & MENG , Q. R. (eds) Mesozoic Sub-Continental Lithospheric Thinning Under Eastern Asia. Geological Society, London, Special Publications, 280, 293–302. DOI: 10.1144/SP280.15 0305-8719/07/$15 # The Geological Society of London 2007.

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Fig 1. (a) The eastern part of the North China Craton; (b) distribution of gold deposits. UHPM belt, ultrahigh-pressure metamorphic belt.

High-resolution seismic tomography Regional seismic tomographic studies mostly have a limited spatial resolution, because of the use of limited datasets of P-wave arrival times and a small number of seismic stations (Liu & Chang 2001; Huang & Zhao 2004; Liu et al. 2005). Low spatial resolution makes it difficult to interpret velocity tomograms. In this study, we gathered abundant datasets that are available to accommodate the high-resolution requirements. We collected the seismic data from 251 stations in the area defined by 1108E –1248E longitude and 348N– 428N latitude, including a total of 8399 earthquakes that occurred from 1993 to 2004. The magnitude of each earthquake is greater than 1.0 ML on the Richter scale. Figure 2 shows the distributions of seismic stations and of regional earthquake epicentres used for this study. The abundant seismic data plus the state-of-the art tomographic methods allowed us to generate the high-resolution tomograms that are needed for understanding the relationship between the velocity variation and crustal heterogeneity. In seismic travel-time tomography, we set up an initial velocity model (Table 1) based on published results of the crustal velocity distribution in the North China Craton and adjacent regions (Liu & Chang 2001; Huang & Zhao 2004; Liu et al. 2005). We adopted a shortest ray path algorithm for ray tracing, and thus have not assumed any a priori discontinuities in the velocity model, such as the Moho discontinuity as a transmitting interface in some conventional ray-tracing methods

(Dijkstra 1959; Moser 1991; Wang & Chang 2000; Chang et al. 2002). The Moho or other boundaries can be identified visually from the resulting tomography images based on velocity contrasts. We parameterized a 3D Earth model as a number of small cubic cells and assigned the velocity to all nodes (cell corners). We defined the travel-time between two connected nodes as their Euclidian distance multiplied by the average slowness (the inverse of velocity). We obtained traveltimes from an epicentre to a station by summing the travel-time from the epicentre to every node until reaching the node corresponding to the station, and selected the shortest ray path based on Fermat’s principle. The inverse problem is a sparse system of linear equations, containing the unknown velocity model and hypocentre parameters. We solved this sparse system using an iteratively damped least-squares algorithm (Bellman 1958; Brown & Dennis 1972; Paige & Saunders 1982; Anderson & Kak 1984). In this study, we used a total of 65 034 P-wave arrival times in the inversion. We defined the cubic cell size as about 29 km  37 km (200  200 ) in the horizontal direction and 4 km in the vertical direction, and thus had 11 088 unknown seismic P-wave velocities to solve. The number of arrival times is five times the unknown velocity nodes, and the sparse linear-equation system is an over-determined problem. To estimate the resolution of our tomographic results, we conducted a series of chequerboard tests (Humphreys & Clayton 1988; Zhao &

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Fig 2. Distribution of the stations (a) and of the earthquake epicentres (b) used in this study. In (a), the red filed triangle shows the position of a station; 93 stations belong to the Huabei Seismic Network, 107 to the Capital Seismic Network, 15 to the Liaoning Seismic Network, and 36 to the Shandong Seismic Network. The total number of stations used in this study is 251. In (b), the red filled circle shows the position of an earthquake epicentre; 3278 events were recorded by the Huabei Seismic Network, 2793 by the Capital Seismic Network, 1950 by the Liaoning Seismic Network, and 378 by the Shandong Seismic Network. A total of 8399 earthquake events were used in this study.

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Table 1. Velocity values in the initial model Depth (km) V0 (km s21)

0–7 4.8

7–17 5.7

17 – 22 6.5

22 – 35 7.6

35 – 40 8.1

Fig 3. Resolution analysis using a chequerboard test for P-wave velocity at depths of 4 km (a), 16 km (b) and 26 km (c).

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Hasegawa 1993; Chang 1996; Liu et al. 2005). We assigned a distinct pattern of velocity perturbations to the cells of the 3D model with a maximum amplitude of +4% alternating perturbation over a background velocity of 6 km s21 for P-wave velocity, and then generated a set of synthetic arrival times by using the actual epicentres and network geometry. We then inverted this synthetic dataset and checked the inversion result visually to identify the resolvability. Figure 3 shows inverted P-wave velocities at depths of 4 km, 16 km and 26 km, on which apparently a distinct perturbation pattern is well reconstructed for the crust. At a depth of 4 km, we have a good resolution in Beijing, Tianjin, Tangshan, Shijiazhuang, and the southern areas of Huhehaote, Zhangjiahou and Chengde where a great number of earthquakes and stations are located, but a low resolution in other areas, as expected. At depths of 16 km and 26 km, we have a good resolution in most regions except Bohai Bay, as a result of a poor station distribution. Bohai Bay is the area surrounded anticlockwise by Dalian, Yingkou, Tangshan, Tianjin and Weihai in Figure 2.

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For real seismic data tomography in the North China Craton, we show reconstructed 3D velocity images in Figures 4–7 as follows: three horizontal slices are at depths of 4 km, 16 km and 26 km and one vertical section is between 1108 and 1248 longitude and 348 and 428 latitude. In these figures, red represents low-velocity zones and blue indicates high-velocity zones, as indicated by the colour bars.

Upper crust and near-surface velocity image Figure 4 is a velocity image at a depth of 4 km that displays the typical lateral heterogeneity of seismic velocity in the upper crust. From Chengde through Zhangjiakou to Huhehaote, an east –west-trending structure with a blue high-velocity perturbation coincides with the Apophysis Belt, situated along the northern boundary of North China Craton. This Apophysis Belt was uplifted for long periods in geological history, underwent magmatic activity during the early Mesozoic (Yanshanian), and late and early Palaeozoic periods (Cheng 1994), and is related to the distribution of gold deposits shown in Figure 1b. Geochemical evidence suggests that

Fig 4. A velocity image slice at a depth of 4 km. Red represents low-velocity zones, and blue high-velocity zones.

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Fig 5. A velocity image slice at a depth of 16 km. Red represents low-velocity zones, and blue high velocity zones.

the ore-forming materials for the gold deposits were generated during Yanshanian time (Chen et al. 1998; Zhai & Yang 2001; Nie et al. 2004). On the southern side of the Apophysis Belt, the velocity anomalies trend NE–SW, coinciding with the regional tectonic structure. The area from 1158E to 1188E outlines part of the North China Plain fault basin, within which the Beijing fault depression shows a lower –velocity perturbation. The area between Beijing and Taiyuan, occupied by the Taihangshan mountain belt, has highvelocity anomalies, and the northern part of Taiyuan between 1138E and 1158E and 398N and 40.58N, which belongs to an early–middle Proterozoic depression, has low-velocity anomalies. This analysis suggests that high-velocity anomalies in Figure 4 are associated with magmatic activity in the Mesozoic, and that the junction between high and low velocities and the NE– SW and east –west fault zones coincides with many mineralization belts.

Middle- and lower-crust velocity images The mid-crustal seismic velocity image at 16 km depth (Fig. 5) is different from the image for the

upper crust. The velocity distribution does not completely coincide with the near-surface structure, and shows strong lateral heterogeneity. The eastern end of the northern margin Apophysis Belt has a lowvelocity perturbation that extends eastwards to Chengde and farther to Yingkou. The southern North China Plain depression has a low-velocity perturbation oriented in a NE–SW direction. The belt from Zhangjiakou through Shijiazhuang to Zhengzhou shows high-velocity anomalies, which almost coincide with the early Proterozoic, north – south-trending, Trans-North China orogen (Zhao et al. 2004). Thus the velocity image evidence is compatible with the fact that the Trans-North China orogen has a deep crystallization basement and metamorphosed volcanic rocks that could be the source of ore-forming material in and around the North China Craton. The Apophysis Belt also coincides with an active earthquake zone. Accordingly, the heterogeneity of the middle crust appears to be an indication of its orogenic activity and instability in crustal history. The velocity image at 26 km depth (Fig. 6) shows a widespread low-velocity perturbation. The low-velocity distribution provides evidence that the lower crust contains high-temperature

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Fig 6. A velocity image slice at a depth of 26 km. Red represents low-velocity zones, and blue high-velocity zones.

Fig 7. A velocity profile along latitude 408N. The horizontal axis indicates longitudes from 1108 to 1248E, and the vertical axis depths in kilometres. Red represents low-velocity zones, and blue high-velocity zones.

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material that is known to be the result of widespread, late Mesozoic basaltic magmatism in the North China Craton (Zhao & Hasegawa 1993; Menzies & Xu 1998; Fan et al. 2000). The high temperature of the lower crust provides feasible conditions for large-scale ore-forming processes in the North China Craton.

Velocity image in vertical section Figure 7 is a vertical section of a velocity image at latitude 408N. In the area from longitude 1178E to 1208E low-velocity perturbations near the surface correspond to the North China sedimentary basin. From 1148E to 1178E a high-velocity perturbation exists in the upper crust above 17 km depth, which corresponds to the position of the TransNorth China orogen; this can be interpreted as the crystalline basement of the Trans-North China orogen, which reaches 17 km depth and trends east– west in the section direction. The area west of latitude 1158E corresponds to the northern part of the Shanxi Uplift where Archaean and lower Proterozoic crystalline basement crops out. Frequent volcanic activity took place in this area during the Mesozoic, especially near latitude 408N, causing widespread lower – velocity perturbations in the upper crust. The area from 1218E to the east corresponds to the eastern end of the Apophysis Belt, the western boundary of which is the Tanlu fault. The position of the Tanlu fault can be identified by the velocity difference in the image. On the eastern, uplift side of the Tanlu fault there are gold deposits that are related to the activity on the fault. The seismic velocity shown in this image is related to crustal temperature and the distribution of basaltic and metamorphosed rocks. Along this profile at c. 20 km depth there is a horizontally uniform lower-velocity zone that has a thickness of less than c. 5 km; it was also demonstrated by Huang & Zhao (2004). The different velocities for the upper and lower crust indicate that the temperature in the lower crust is higher than that in the upper crust. This clearly reflects the different material components and geophysical characteristics of the upper and lower crust.

Discussion We have reconstructed 3D velocity images of the North China Craton using seismic tomography that demonstrate the lateral heterogeneity of the crust. The reconstructed velocity images indicate that there are east –west- and NE–SW-trending structures at shallow depths that correspond to structures on the surface, and that the distribution

of Cretaceous gold deposits and epicentres reflects the strikes of known east –west- and NE –SWtrending faults. Comparing the location of gold deposits and epicentre distribution with the velocity images we find that most gold deposits and earthquakes are located on the tectonic boundaries of the North China Craton, which coincide with highvelocity anomalies. These boundaries underwent major tectonic and igneous activity in the Mesozoic related to the well-established removal, thinning or delamination of sub-continental lithosphere, and thus the velocity distributions provide independent evidence for such delamination.

Conclusions The 3D velocity images reconstructed from the most recent seismic data display crustal heterogeneities of the North China Craton. The P-wave velocity distributions in 3D crust demonstrate the following features. (1) The collision and amalgamation of the North China Craton with contiguous continental blocks are reflected by east– west- and NNE – SSW-trending high- and low-velocity boundaries. (2) High-velocity anomalies coincide with prominent late Mesozoic magmatic and mineralization belts along the boundary faults of the North China Craton. (3) The crystalline basement of the North China Craton reaches a depth of 17 km. (4) Widespread, low-velocity anomalies in the lower crust are indicative of large-scale mantle upwelling, and are consistent with the idea that considerable sub-continental lithosphere removal took place in the Cretaceous, a model that is independently supported by many papers in this volume. This research was supported by the Natural Science Foundation of China (Grants 40235055 and 40230056).

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Fluid evolution and large-scale gold metallogeny during Mesozoic tectonic transition in the Jiaodong Peninsula, eastern China H.-R. FAN, F.-F. HU, J.-H. YANG & M.-G. ZHAI Key Laboratory of Mineral Resources, Institute of Geology and Geophysics, Chinese Academy of Sciences, PO Box 9825, Beijing 100029, China, (e-mail: [email protected]) Abstract: The Jiaodong Peninsula or eastern Shandong Province, the most important gold producing in region China, is located in the southeastern margin of the North China Craton. The gold deposits in the Jiaodong Peninsula are divided into three gold belts, from west to east, the Zhaoyuan– Laizhou, Penglai–Qixia and Muping–Rushan belts. The deposits occur as goldbearing quartz veins and disseminated- and stockwork-style ores adjacent to fault zones. Most of the gold deposits can be classified in four stages: stage I, quartz– (minor) pyrite; stage II, pyrite–quartz– gold; stage III, quartz– base metal sulphide minerals; stage IV, quartz– carbonate. Ar–Ar ages, Rb–Sr isochrons, and hydrothermal zircon sensitive high-resolution ion microprobe U– Pb ages obtained from these deposits suggest a gold mineralization time of 120 + 10 Ma. The Sr –Nd isotopic compositions of pyrites and the associated rocks suggest that the ore-forming materials were probably derived from a mixed source. Fluid inclusion studies show that oreforming fluids of gold deposits are consistent throughout the Jiaodong Peninsula, with similar mineralizing temperature and pressure conditions. Ore-forming fluids are characterized by H2O– CO2 – NaCl + CH4. The optimal mineralizing temperature and pressure ranges are 170– 335 8C and 0.7–2.5 kbar. Oxygen and hydrogen isotope data show that ore fluids are of magmatic origin. Gold deposits in the Jiaodong Peninsula formed in the same mineralizing– geodynamic conditions, and are related to the Mesozoic tectonic transition in the eastern North China Craton. Gold metallogeny is only one expression of the Mesozoic tectonic transition.

The Jiaodong gold province in the Jiaodong Peninsula of Eastern China (Fig. 1) is currently the most important gold producing region in China. Seven worldclass gold deposits (.100 t gold), eight large gold deposits (20–100 t gold) and more than 100 medium to small gold deposits (,20 t gold) have been discovered in the Peninsula during the past three decades, accounting for about 25% of Chinese gold reserves (Zhou & Lu 2000; Fan et al. 2003). The genesis of gold deposits in the Peninsula is a popular research subject for economic geologists. Some researchers (Shen et al. 1994; Wang 1996) considered them to be the result of Precambrian metamorphic hydrothermal fluid. However, it is widely considered now that the formation of the gold deposits is closely related to Mesozoic tectonic–magmatic hydrothermal activity or collisional orogenesis (Wang et al. 1998; Zhai et al. 2001, 2002, 2004; Qiu et al. 2002; Fan et al. 2003; Mao et al. 2003, 2005; Yang et al. 2003; Chen et al. 2004, 2005). Previous studies on ore-forming fluids of gold deposits in the Peninsula were mostly published in China and mainly focused on a single deposit (Xu et al. 1996; Zhai et al. 1996; Lu et al. 1999; Shen et al. 2000; Qiu et al. 2002; Fan et al. 2003; Hu et al. 2005), and have not provided a regional-scale model of ore-forming fluids. Our study builds on fluid inclusion data for the most important gold

deposits in the Peninsula, and discusses ore-forming fluid evolution and gold lode genesis during the Mesozoic tectonic transition in the Jiaodong Peninsula.

Geological setting Regional and mine geology The Jiaodong gold province occurs along the southeastern margin of the North China Craton, which is dominated by Archaean rock units. It is bounded by the north–south- to NE–SW-trending Tanlu fault zone to the west and by the Sulu ultrahigh-pressure metamorphic (UHPM) belt to the SE (Fig. 1). Supracrustal rocks in the Jiaodong Peninsula comprise both metamorphosed Precambrian sequences and Mesozoic volcanic rocks and intrusions. The Precambrian sequences are composed of basement rocks of the Late Archaean Jiaodong Group and the Palaeoproterozoic Jinshan and Fengzishan Groups, which consist of mafic to felsic volcanic and sedimentary rocks metamorphosed to amphibolite or granulite facies. Plutonic rocks, which intruded into the Precambrian basement in the Peninsula, include the widely distributed Mesozoic Linglong, Guojialing, Luanjiahe, Aishan, Kunyushan and Weideshan granitoids.

From: ZHAI , M.-G., WINDLEY , B. F., KUSKY , T. M. & MENG , Q. R. (eds) Mesozoic Sub-Continental Lithospheric Thinning Under Eastern Asia. Geological Society, London, Special Publications, 280, 303–316. DOI: 10.1144/SP280.16 0305-8719/07/$15 # The Geological Society of London 2007.

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Fig. 1. Simplified geological map of the Jiaodong Peninsula showing location of the major gold deposits (after Qiu et al. 2002; Fan et al. 2003). Different symbol sizes indicate gold resources sizes: large symbol, Au .50 t; small symbol, Au ,50 t.

Gold deposits in the Peninsula can be divided into three mineralized belts from west to east, namely the Zhaoyuan–Laizhou belt, Penglai–Qixia belt and Muping–Rushan belt (Fig. 1). The belts are separated by Jurassic to Cretaceous volcanic– sedimentary basins. Most gold deposits are distributed to the west of the Mishan fault, although sporadic small-scale gold deposits have been reported in the granitoids of Wendeng and Weihai Cities (Fig. 1). There are two main types of gold deposits in the Peninsula, namely quartz vein-style (Linglong type) and fault-zone hosted disseminated and stockworkstyle (Jiaojia type) (Qiu et al. 2002; Fan et al. 2003). Linglong-type quartz vein-style gold deposits are represented by the Linglong, Jiuqu, Denggezhuang and Jinqingding gold mines (Fig. 1). Gold mineralization is developed in quartz veins, and typically hosted in the second- or third-order faults cutting Mesozoic granitoids. This mineralization occurs as single or multiple relatively continuous quartz veins that may be several hundred metres to .1 km in strike, range from a few centimetres to a few metres in width, and extend for at least a few hundred metres down dip. This style of gold mineralization is typical in the Linglong gold mine, where more than 100 auriferous quartz veins are hosted in the second- or third-order faults. The lodes are rarely zoned, from a central massive sulphide zone as thick as 1 m and consisting of pyrite with minor chalcopyrite, galena and sphalerite, to

marginal zones of quartz–pyrite veins. The quartz veins are translucent to milky and grey, and appear to fill pre-existing faults, as they are commonly bounded by fault gouge or thin zones of quartz–sericite schist. The veins may contain fragments of the wall-rocks and also rarely cut the foliation of the bounding schist or the fault gouge, suggesting that at least some of the host faults were formed prior to the quartz veins. Most of the quartz veins, and particularly the zones of massive sulphide minerals within the veins, have been fractured, brecciated or boudinaged. The veins occur as lenticular discrete bodies bounded by the fault gouge, although, in rare cases, undeformed combtextured quartz veins are present. Wall-rock alterations adjacent to the quartz veins commonly include silicification, sericitization, sulphidation, and potassic alteration. Degrees of wall-rock alteration in the different gold mines or quartz veins are different. Transitional styles between Linglong-type and Jiaojia-type deposits are found in the deep parts of some mines. Jiaojia-type fault-zone hosted disseminated and stockwork-style gold deposits are represented by the Sanshandao, Jiaojia, Xincheng, and Yingezhuang gold mines (Fig. 1). These disseminatedand stockwork-style deposits occur along the firstorder regional faults, which are gently dipping and are surrounded by broad alteration haloes. The faults commonly crop out as cataclastic

MESOZOIC FLUIDS AND GOLD, EAST CHINA

deformation zones that are a few hundred metres wide and, although sometimes described as shear zones, they show few or no ductile features. They are characterized by extensive wall-rock alteration or superimposed late brittle deformation. In the fault zones with Jiaojia-type mineralization, a main straight and smooth fault plan is well developed, and gold ore bodies are mainly located in the footwall of the main fault plane. Outward from the main fault plane, the following alteration zones are very common within the granitoid hosted Jiaojia-type deposits. (1) A fault gouge zone may be up to 50 cm thick, and consists of grey, white, or black clay materials (commonly .80%) with variable-sized, round fragments of wall-rock granitoids. Industrial gold mineralization is rarely developed, with uncommon gold mineralizing fragments existing locally. (2) A pyrite –sericite zone may be up to several metres to .10 m thick in some places. It comprises quartz, sericite and disseminated pyrite, with minor calcite, and is the product of extensive wall-rock alteration. (3) A pyrite –sericite granitoid zone may be up to 50 m thick, with disseminated pyrite and pyrite– quartz stockworks. It is composed of silicified, sericitized and pyritic granitoid. (4) An outer reddish (K-feldspar) alteration zone extends for up to several hundred metres, and is characterized by Kfeldspar and sericite in cataclastically deformed granitoids. Fine quartz veins with pyrite are developed in the reddish zone and industrial gold ores may be formed in crowded pyrite– quartz stockworks. So-called interstratified glide breccia-style (Pengjiakuang type) gold deposits in the Peninsula described by some workers (Shen et al. 1998) are a special example of the Jiaojia type disseminatedand stockwork-style mineralization. These ores are located in the margin of the Cretaceous Jiaolai basin, and hosted in the interstratified low-angle glide faults between basement and cover rocks. Wall-rocks include Precambrian metamorphic rocks of the Jinshan and Jiaodong Groups, and Mesozoic granitoids or dykes. The wall-rocks have undergone extensive ductile, glide or fragment deformation with silicification, sericitization, carbonation and pyritization.

Ore-forming paragenesis and major ore minerals Previous researchers have proposed different mineralization epochs and stages for gold deposits in the Jiaodong Peninsula. Chen et al. (1989) suggested six mineralization stages, Yao et al. (1990) proposed five mineralization stages, and Luo & Miao (2002) four mineralization stages. Gold lodes in the Jiaodong

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Peninsula are characterized by superimposition and enrichment that occurred in multiple epochs and stages. Multiple structure–hydrothermal fluid activity controlled gold mineralization. Based on the mineralogical, textural and crosscutting relationships observed in the field and under the microscope, four stages of hydrothermal mineral formation have been identified; from the oldest to the youngest, these are: stage I, quartz–(minor) pyrite; stage II, pyrite–quartz–gold; stage III, quartz–base metal sulphide minerals; stage IV, quartz–carbonate. Gold mineralization mainly took place in stages II and III. Stage I, quartz–(minor) pyrite, is pre-gold paragenesis, and consists of coarse-grained quartz (95%) and minor pyrite. Quartz is milky-white in colour with a greasy lustre and subhedral habit. Some wall-rock alteration fragments such as sericite exist in the quartz aggregations. Pyrite is light yellow –white in colour with a strong metallic lustre, and occurs as euhedral cubes. Fractures are well developed in the pyrite, and later stage minerals, such as galena, sphalerite, chalcopyrite, electrum and native gold, may occur as infillings. Stage II, pyrite –quartz–gold, is the main gold mineralization epoch. The mineral assemblage is dominantly composed of quartz, pyrite, electrum and native gold. Quartz is fine-grained and granular, with a smoky-grey colour and vitreous lustre. Pyrite, exhibiting a light yellow colour with a weak metallic lustre, generally occurs as euhedral pentagonal or cubic shapes or fine-grained subhedral aggregates. Fine-grained gold occurs in quartz or sulphide minerals, particularly in pyrite. Some arsenopyrite occurs as fine-grained subhedral and anhedral aggregates in the Sanshandao and Linglong deposits. The mineral assemblage in stage III, quartz–base metal sulphide minerals, is more complex and also represents a major gold mineralization epoch. Ore minerals are dominantly composed of pyrite, chalcopyrite, sphalerite, pyrrhotite, native gold, and electrum. Gangue minerals are mainly quartz, with minor sericite and calcite. This paragenesis forms veinlet and disseminated mineralization. Quartz is smoky-grey in colour, and pyrite occurs as fine-grained subhedral and anhedral aggregates. The other sulphide minerals occur as fine-grained anhedral aggregates, typically filling microcracks in the early pyrite and quartz. Gold mainly fills cracks within pyrite and arsenopyrite, and less commonly within sphalerite and galena. Stage IV, quartz–carbonate, is the latest epoch. The mineral assemblage mainly consists of quartz and carbonate minerals with minor fine pyrite, occurring as veinlets filling late fractures. Carbonates include calcite, ankerite and siderite. Calcite is typically anhedral with well-developed twins. Fine-grained subhedral quartz is distributed in the

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calcite aggregates. This stage post-dates the gold mineralization because it cuts the mineralized veins. There are differences in ore assemblage, type, structure and morphology for the four mineralization stages mentioned above, which can be used to distinguish the stages in the field and laboratory. As multiple phases of mineralization have occurred in the Jiaodong gold province, in particular the superimposition and enrichment of the stage II and III mineralization, these add to the difficulty in distinguishing between the stages. As quartz in both stages II and III is smoky-grey in colour and cannot easily be distinguished in the field and laboratory, fluid inclusions in quartz of stages II and III are put into the same generation (gold stage) in the following discussion.

Fluid inclusions Microthermometric measurements of fluid inclusions were carried out on a Linkam THMS 600 heating– freezing stage at the State Key Laboratory of Lithosphere Evolution and Key Laboratory of Mineral Resources, Institute of Geology and Geophysics, Chinese Academy of Sciences (IGGCAS). Temperature was calibrated using a set of synthetic inclusions from Fluid Inc., USA. The minimum measurement temperature was 2195 8C. The temperature measurements can be reproduced within +0.5 to +0.2 8C in the range of 2120 to 0 8C, and within +0.2 to +2 8C in the range of 0–550 8C. Most measurements were made at a heating rate of 1 8C min21 at lower temperature (,30 8C), and at a heating rate of 3 8C min21 at higher temperature (.100 8C). The heating rate near the temperature of phase transformation was ,0.2 8C min21 for measurements. Raman microspectrometer analyses have been made with a Ranishaw MK1-2000 laser Raman microspectrometer at IGGCAS, using a 514 nm Ar-ion laser as the source of excitation. The MacFlincor software (Brown & Hagemann 1995) and a computer program by Bakker (1997) to calculate salinity from clathrate melting temperature were chosen to handle the measurement data. Molar volumes of CO2-bearing inclusions were calculated from the diagram of Thie´ry et al. (1994). The salinity of aqueous inclusions was based on the temperature of last ice melting during heating runs, and was obtained using the equation of Bodnar (1993).

Fluid inclusion petrography Samples for fluid inclusion study were collected from representative gold deposits in the Jiaodong Peninsula; from west to east, these are the

Sanshandao, Yinggezhuang, Linglong, Daliuhang, Denggezhuang and Jinqingding deposits (Fig. 1). Of these, the Sanshandao and Yinggezhuang deposits are disseminated and stockwork-style (Jiaojia type), the Daliuhang deposit is transitional between disseminated- and stockwork-style and quartz vein-style, and the Linglong, Denggezhuang and Jinqingding deposits are quartz vein-style (Linglong type). Based on thermometric measurements and Raman microspectrometer analyses, three types of fluid inclusions have been distinguished from quartz in gold lodes: A-type H2O–CO2 inclusions; B-type CO2 –H2O + CH4 inclusions; C-type aqueous H2O inclusions. A-type H2O–CO2 inclusions have negative crystal or regular shapes. These inclusions, mostly 5–10 mm in size, generally consist of two to three phases (Fig. 2a), an aqueous phase, a liquid CO2 and/or a gas CO2 bubble, at room temperature, showing a VCO2 þ LCO2 of ,55% of whole volume (Fig. 2a). B-type CO2 –H2O + CH4 inclusions (Fig. 2b–f) also show two to three phases, an aqueous phase, a liquid CO2 and/or a gas CO2 bubble, at room temperature. VCO2 þ LCO2 in the inclusions is variable, ranging from 15 to 60%, and may sometimes reach 90% (Fig. 2b). The inclusions have negative crystal or regular shapes, and generally range in diameter from 6 to 15 mm and to rarely to 40 mm (Fig. 2c). C-type aqueous H2O inclusions consist of an aqueous liquid and a minor vapour phase (VH2O þ LH2O), and have a gas–liquid content of 5–30%. The inclusions are tabular, elliptical, or have irregular shapes (Fig. 2g and h) with a variable size of ,1–20 mm. Fluid inclusion type combinations in quartz and carbonate minerals formed in different stages vary. Milky-white quartz in stage I was formed earlier, and hosted the late fluid inclusion types. Primary fluid inclusions in milky-white quartz are H2O–CO2 inclusions (Fig. 2a) and aqueous H2O inclusions of small size (commonly ,4 mm) (Fig. 2g). Differentiation H2O–CO2 and aqueous H2O inclusions at room temperature should be based on microthermometric or laser Raman measurements. Quartz formed at stage I has a milky-white colour because it hosts a large quantity of small fluid inclusions, both primary inclusions and late inclusions trapped by extensive micro-cracking and healing of quartz (Boiron et al. 1996). B-type CO2 –H2O + CH4 inclusions, occurring along planar arrays in healed microfractures in stage I quartz, are obviously secondary. B-type inclusions in stage II and III quartz occur at random and show a primary or pseudosecondary origin. In some deposits such as the Sanshandao and Linglong deposits, CO2 – H2O + CH4 inclusions in stage III quartz have more variable CO2 and H2O, showing the coexistence of several H2O–CO2 inclusions (Fig. 2f) with similar

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Fig. 2. Photomicrographs of typical fluid inclusions of gold deposits in the Jiaodong Peninsula showing: (a) H2O– CO2 fluid inclusions in a pre-gold quartz vein; (b) isolated CO2 –H2O + CH4 inclusions in a gold-stage quartz vein; (c) isolated three-phase CO2 –H2O + CH4 inclusions in a gold-stage quartz vein; (d) three-phase CO2 –H2O + CH4 inclusions in a gold-stage quartz vein; (e) primary CO2 –H2O + CH4 inclusions around pyrite in a gold-stage quartz vein; (f) two- and three-phase CO2 –H2O + CH4 inclusions with varying phase ratios in a gold-stage quartz vein; (g) primary aqueous fluid inclusions in a pre-gold quartz vein; (h) secondary aqueous fluid inclusions within a gold-stage quartz vein. All inclusions are in quartz.

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homogenization temperatures. This feature indicates that ‘boiling’ or immiscibility may have occurred when the fluids were trapped (Roedder 1984; Shepherd et al. 1985; Lu et al. 2004). Aqueous H2O inclusions in the gold stage (stages II and III) occur as secondary trails in microfractures cutting early quartz grains (Fig. 2h) or as random inclusions in quartz and calcite in stage IV veins. These inclusions are interpreted to contain a fluid trapped late in the hydrothermal history.

Microthermometric results The microthermometric data for all types of fluid inclusions are summarized in Table 1 and discussed below. Pre-gold H2O–CO2 inclusions and aqueous H2O inclusions. The H2O–CO2 fluid inclusions freeze to a solid phase below 295 8C. During heating runs, melting of the carbonic phase (TmCO2) occurs either at the CO2 triple point of 256.6 8C, or over a small interval with depressed melting temperatures between 258.1 and 256.6 8C (Table 1). These measurements indicate that the carbonic phase in these inclusions is nearly pure CO2 that probably contains minor amounts of additional gas species, probably CH4 and/or N2 (Burruss 1981). Melting of the CO2 clathrate (Tmcl) in the presence of CO2 liquid occurs between 4.0 and 10.2 8C (Table 1), below and just above the invariant point of a pure CO2 clathrate (10 8C; Hollister & Burruss 1976). Partial homogenization (ThCO2) of CO2 liquid þ CO2 vapour to liquid CO2, or less commonly to vapour, occurs between 7.4 and 29.8 8C. During heating runs, more than one-half of the studied H2O–CO2 inclusions decrepitated prior to final homogenization, at temperatures from 215 to 250 8C. Temperatures of total homogenization to liquid, obtained mainly from inclusions with lower CO2 contents and smaller diameters, range from 250 to 403 8C (Table 1). As Raman peaks were not detected for N2 and H2S in these inclusions, these species are assumed to be very minor. The molar fractions of CO2 and CH4 (xCO2 and xCH4) in the carbonic phase calculated using the method of Bakker (1997) and MacFlincor software (Brown & Hagemann 1995) are 0.5 (commonly 0.3) to 0.1 and ,0.08, respectively. The CO2 densities of the carbonic phase are between 0.80 and 0.95 g cm23, and densities of the bulk inclusion range from 0.79 to 1.03 g cm23. Calculated salinities of the aqueous phase in these inclusions range from 10.8 to 0.0 wt% NaCl equiv. As primary aqueous H2O inclusions are very small in stage I quartz, commonly ,4 mm, it is not easy to

make microthermometric measurements. Homogenization temperatures, 243–286 8C, of primary aqueous H2O inclusions were obtained only from barren milky-white quartz from the Linglong deposit. Melting of ice (Tmice) of these inclusions ranges between 24.3 and 21.5 8C (Table 1), corresponding to salinities from 6.9 to 2.6 wt% NaCl equiv. The cross-cutting relations between H2O– CO2 inclusions and early H2O inclusions have not been observed in stage I milky-white quartz, and conclusions on the generation for these two inclusions are uncertain. Gold stage (stages II and III) CO2 –H2O + CH4 inclusions. The TmCO2 during heating runs of the CO2 – H2O + CH4 inclusions is from 265.6 to 256.6 8C (Table 1), and thus suggests the presence of significant amounts of other non-aqueous phases such as CH4, in addition to CO2, in many inclusions. The presence of CH4 was verified by laser Raman microspectrometry (Hu et al. 2005). The Tmcl measurements are between 3.6 and 13.5 8C for these inclusions. Partial homogenization of CO2 – CH4 liquid and CO2 –CH4 vapour, consistently to a liquid, occurs over a wide range from –8.3 to 30.1 8C. The inclusions show a range in final homogenization temperatures to liquid from 170 to 335 8C (Table 1). It is important to note that these data are mainly from inclusions with relatively small CO2 + CH4 bubbles (,40 vol.%), because almost all the inclusions with greater volumes of the carbonic phase decrepitated prior to homogenization. The presence of significant CH4 in the inclusions results in increased clathrate melting temperatures up to more than 10 8C (Lu et al. 2004). The calculated mole proportions of CO2 and CH4 in the carbonic phases are 0.99 –0.74 and 0.26 –0.01, respectively, assuming that the carbonic phases are mainly CO2 and CH4. Using MacFlincor software (Brown & Hagemann 1995) and the computer programs of Bakker (1997), calculated xCO2 and xCH4 in the bulk inclusion are 0.04 –0.50 and ,0.01 –0.14, respectively. The calculated densities of the carbonic phase and bulk inclusion are 0.62 – 0.98 g cm23 and 0.82 –1.01 g cm23, respectively. The Tm data (Table 1) indicate salinities of ,11 to 0 wt% NaCl equiv., if the CO2 – H2O –NaCl system is used as a reference (Diamond 1992). In the Sanshandao and Denggezhuang samples, some inclusions trapped fluids with higher CH4 content (Fan et al. 2003; Hu et al. 2005). Post gold (stage IV) aqueous H2O fluid inclusions. The aqueous H2O inclusions show final homogenization to liquid at temperatures between 96 and 228 8C, and melting of ice in the range of 20.1 to 25.0 8C (Table 1). The melting temperatures correspond to salinities from 0.2 to 9.9 wt% NaCl equiv. (Bodnar 1993).

Fluid inclusion type

H2O – CO2 inclusion

CO2 – H2O + CH4 inclusion

Aqueous inclusion

Ore stage

Stage I (pre-gold stage)

Stages II and III (gold stage)

Stage IV (post-gold stage)



(

Linglong Yingezhuang Daliuhang Denggrzhuang Jinqingding Sanshandao Linglong Yingezhuang Daliuhang Denggrzhuang Jinqingding Sanshandao Linglong Yingezhuang Daliuhang Denggrzhuang Jinqingding

(Sanshandao

Sample location Tmcl (8C) 5.4– 10.2 5.6– 8.9 5.5– 8.9 6.5– 9.1 5.8– 8.7 4.0– 7.9 6.4– 13.5 4.4– 12.2 5.1– 9.4 6.3– 8.7 4.3– 9.0 3.6– 9.1

TmCO2 (8C) 257.5 to 256.7 257.1 to 256.6 258.1 to 256.7 256.8 to 256.6 256.9 to 256.6 256.7 to 256.6 265.6 to 256.8 260.2 to 256.7 259.0 to 256.8 257.1 to 256.6 262.4 to 256.8 258.1 to 56.6

Table 1. Microthermometric data for fluid inclusions of gold deposits in the Jiaodong Peninsula

7.4– 22.4 14.1– 22.3 20.1– 29.8 17.2– 28.3 15.6– 27.5 7.5– 28.1 28.3– 26.8 1.6– 30.1 14.2– 29.8 17.2– 28.2 7.2– 29.3 4.8– 31.0

ThCO2 (8C)

25.0 26.5 25.8 24.3 25.5 22.8

to to to to to to

20.4 21.2 20.5 20.3 20.9 20.1

Tmice (8C)

267 – 375 256 – 345 280 – 355 250 – 347 254 – 365 305 – 403 204 – 325 210 – 335 185 – 283 189 – 276 195 – 317 170 – 324 143 – 228 120 – 215 110 – 185 152 – 195 156 – 219 96 – 188

Thtot (8C)

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Hydrothermal fluid evolution and gold mineralization Ore-forming fluid evolution The pre-gold H2O–CO2 fluid inclusions occur isolated in samples of the early milky-white quartz and are assumed to have been trapped during crystallization of the earliest quartz. Fluid inclusions associated with gold mineralization contain mostly CO2 –H2O + CH4 fluids, and occur isolated, and sometimes in healed fractures, particularly in smokygrey quartz with variable CO2 and CH4 contents. Figure 3 shows the homogenization (ThCO2) v. melting temperature (TmCO2) of pre-gold H2O–CO2 and gold-stage CO2 –H2O + CH4 inclusions in the Sanshandao gold deposit. The molar volumes of H2O–CO2 and CO2 –H2O + CH4 inclusions are 50–60 cm3 mole21 and 55–70 cm3 mol21, respectively. There are similar changes of molar volumes of H2O–CO2 and CO2 –H2O + CH4 inclusions in other deposits of the Jiaodong Peninsula. Variable CO2 molar fractions are especially common for the gold-stage CO2 – H2O + CH4 inclusions at the Sanshandao deposit (Figs 3 and 4). This can be caused by several factors, such as immiscibility and fluid –wall-rock interactions (e.g. Ramboz et al. 1982; Craw et al. 1993; Nabelek &

Ternes 1997). Heterogeneous trapping of the immiscible fluids would result in various phase ratios and two different modes of final homogenization (L, V) (Ramboz et al. 1982). Such variability was observed only in the CO2 – H2O + CH4 inclusions of some of the studied sections. If immiscibility did occur at Sanshandao, then it was only of local importance. The existence of CO2-bearing inclusions in the pre-gold and gold-stage quartz in the Jiaodong gold province provides the possibility to estimate trapping pressures of the inclusions by constructing isochores from microthermetric and fluid composition data (Shepherd et al. 1985). Using MacFlincor software (Brown & Hagemann 1995), estimated fluid trapping pressures of the pre-gold and gold stage are 1.0–3.5 kbar and 0.7– 2.5 kbar, respectively. The pressure tends to become lower with fluid evolution.

Timing of large-scale gold mineralization In recent years, detailed dating of gold mineralization in the Jiaodong Peninsula has been carried out using the Ar –Ar method on sericite, muscovite, K-feldspar and quartz, the Rb– Sr method on pyrite, and sensitive high-resolution ion microprobe (SHRIMP) on hydrothermal zircon (Table 2).

Fig. 3. Homogenization (ThCO2) v. melting temperature (TmCO2) of H2O– CO2 and CO2 –H2O + CH4 inclusions in the Sanshandao gold deposit of the Jiaodong Peninsula. In the figure, the relations are shown between (1) molar fractions CH4 (XCH4 ¼ 0.05– 0.15), (2) molar volumes (50– 70 cm3 mol21) and various combinations of homogenization temperatures and melting temperatures for fluid inclusions in the CO2 – CH4 system (data from van den Kerkhof 1990; Thie´ry et al. 1994). Tp CO2 is the triple point of pure CO2; Tc CO2 is the critical temperature of pure CO2.

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Fig. 4. Total homogenization temperature (Th) v. calculated equivalent mole fraction CO2 of CO2 –H2O + CH4 inclusions in the Sanshandao gold deposit of the Jiaodong Peninsula. The curves delimit the two-phase regions in the H2O– 6 wt% NaCl system at 0.5 and 1 kbar. l, liquid; g, gas; c, critical; decr, decrepitation (after Bowers & Helgeson 1983).

Ages of different deposits (including the special interstratified glide breccia-style of Jiaojia type) obtained by different methods show no difference. Large-scale gold metallogeny in the Jiaodong Peninsula occurred in the period of 120 + 10 Ma. Gold mineralization occurred in a short period, and formed in the same mineralizing background and from the same tectonic– magma–fluid mineralizing system. Although mineralization stages can be differentiated by their ore-forming paragenesis and alteration relations, these stages are the results of the same metallogeny.

Possible sources of ore-forming fluids and metallogenic materials Previous discussion on sources of metallogenic materials are based on sulphur and carbon stable isotope data. As remarkable stable isotope fractionation often occurred during geological and

geochemical processes, uncertainty in tracing the source of metallogenic materials has risen greatly. Sr and Nd isotopic data can solve this problem. Zhou et al. (2003) analysed the Sr–Nd isotopic composition of ores (pyrite), granites, intermediate– basic dykes and basement metamorphic rocks in the Jiaodong Peninsula. Liu et al. (2003) also measured the Sr –Nd isotopic composition of carbonate minerals from the Jiaojia and Jinqingding gold deposits. The results (Fig. 5) show that the Sr –Nd isotopic compositions of ores are close to those of mantle rocks or mafic dykes in the Peninsula, and also have a relation to Precambrian basement and Mesozoic granites. The gold and related metallogenic materials were derived from a mixed source, and mantle components were involved in ore-forming processes (Zhou et al. 2003). Hydrogen, oxygen, carbon and sulphur stable isotope analyses (Xu et al. 1996; Zhai et al. 1996; Lin & Yin 1998; Mao et al. 2002; Fan et al. 2003; Liu et al. 2003) demonstrate that the initial

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Table 2. Isotope ages determined in the lasts 5 years for gold deposits in the Jiaodong Peninsula Deposit/type Cangshang/ Jiaojia type Jiaojia/ Jiaojia type Xincheng/ Jiaojia type Wangershan/ Jiaojia type Dongji/ Jiaojia type Linglong/ Linglong type Dazhuangzi/ Jiaojia type Penjiakuang/ Jiaojia type Fayunkuang/ Jiaojia type Jinqinding/ Linglong type

Sample

Method

Age (Ma)

Reference

Sericite

Ar –Ar

121.3 + 0.2

Zhang, X. O., et al. (2003)

Sericite, muscovite Sericite, muscovite Sericite, muscovite K-feldspar, quartz Pyrite

Ar –Ar

120.5 + 0.6 to 119.2 + 0.2

Li, J. W., et al. (2003)

Ar –Ar

120.7 + 0.2 to 120.2 + 0.3

Li, J. W., et al. (2003)

Ar –Ar

121.0 + 0.4 to 119.4 + 0.2

Li, H. M., et al. (2003)

Ar –Ar

116.3 + 0.8 1144 + 0.2

Li, J. W., et al. (2003)

Rb –Sr

122.7 + 3.3 to 123.0 + 4.2

Yang & Zhou (2001)

Quartz

Ar –Ar

115.6 + 1

Zhang, L. C., et al. (2003)

Quartz, biotite Pyrite

Ar –Ar

117.5 + 0.3 to 118.4 + 0.3

Zhang, L. C., et al. (2003)

Rb –Sr

128.2 + 7

Zhang, L. C., et al. (2003)

Hydrothermal zircon

SHRIMP U–Pb

117 + 3

ore-forming fluids were composed mainly of magmatic water, and were added by meteoric water during later metallogenic processes. The close temporal relation between the gold deposits and the

Hu et al. (2004)

120–126 Ma mafic dykes (Yang et al. 2003) may indicate that the magmatic water was derived from the degassing of magmas parental to the dykes. Therefore, it is concluded that the ore

Fig. 5. Isotopic compositions of Sr and Nd for pyrites and carbonate minerals in the gold deposits in the Jiaodong Peninsula, compared with those of associated rocks in space and time. Data for carbonate minerals from the Jiaojia and Jinqingding gold deposits are after Liu et al. (2003); basement metamorphic rocks after Yang & Zhou (2001); others after Zhou et al. (2003).

MESOZOIC FLUIDS AND GOLD, EAST CHINA

fluids responsible for mineralization in the Jiaodong gold province were derived from fluids degassed from mafic to intermediate magmas (Fan et al. 2003; Zhou et al. 2003).

Tectonic – magma – fluid geodynamic background The North China Craton, in which the Jiaodong gold province is located, was welded to the eastern part of the Eurasian continent by the Dabie –Sulu collision orogeny in the south and the Xingmeng collision orogeny in the north during the late Palaeozoic and early Mesozoic, and then became part of the NNE–SSW-trending marginal Pacific tectonic domain (Zhai et al. 2003; Kusky et al. 2007). Evolution of continental crust in the region entered a new period of intra-continental orogeny. This intense tectonic inversion marked the end of pre-Mesozoic basin–mountain systems, and established the basic tectonic pattern of the North China Craton from the late Mesozoic to the present. In recent years, Chinese geologists have proven that there exists an intense inversion of tectonic mechanism during the Mesozoic in eastern North China (Zhai et al. 2003). This is reflected by the following features. (1) Tectonically, the east–westtrending structural framework changed to a NE– SW- to ENE–WSW-trending structural framework, and a compressive tectonic regime during the Palaeozoic changed to an extensional regime during the Jurassic–Middle Cretaceous. The basin–mountain system experienced great changes or reconstruction (Meng 2003). (2) Geodynamically, the eastern Central Orogenic Belt in China experienced deep continental subduction, followed by exhumation of ultrahigh-pressure slabs to the surface (Kusky & Li 2003). The regime of collision and amalgamation of different continental blocks changed to an intracontinental tectonic regime (Zhai & Fan 2002). (3) In terms of lithosphere evolution and change in thickness, large-scale mantle upwelling and lithosphere removal took place (Zheng 1999; Gao et al. 2004). The lithosphere thickness was more than 200 km (Chi & Lu 1996) in the Palaeozoic, but thinned to less than 80 km in the Mesozoic (Fan & Hooper 1991; Zhou et al. 2003). (4) Little magmatism took place in the Proterozoic and Palaeozoic, whereas large-scale acidic magmatism began in the Mesozoic (Wu 1985) and reached a peak at 120–110 Ma, indicating strong mantle–crust interaction (Zhang et al. 2002). (5) Geological fluids and large-scale mineralization are widespread in the eastern North China Craton (Zhai et al. 2001; Yang et al. 2003; Chen et al. 2004; Mao et al. 2005). As a direct response to tectonic inversion and lithosphere thinning during the Mesozoic, continental crust, particular lower crust, in eastern North

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China experienced extensive recombination and remelting. Large-scale upwelling of the mantle induced disturbance, rotation, and adjustment of heat flow and density structure, resulting in major remelting of crust and production of acidic magma, material exchange and mixing of mantle– crust, and transport of fluids, finally forming a new magma– fluid– mineralization system (Zhai et al. 2003). These processes induced the intensive and large-scale super-accumulation of metals (gold mineralization) in a short time in the eastern North China Craton. Different types of gold deposits in the Jiaodong Peninsula were formed in the same mineralizing background related to the Mesozoic tectonic inversion; these are only one manifestation of the Mesozoic tectonic inversion. From the available data, lithosphere thinning in the Jiaodong Peninsula was a continuous process. During the late Jurassic (160–140 Ma), crustal source adamellite was formed as large areas, represented by the Linglong, Kunyushan and Luanjiahe granitoids (Luo & Miao 2002; Zhang et al. 2004), indicating a relatively thick crust. In the middle Cretaceous, granodioritic magma from a relatively deep source intruded as stocks, represented by the Guojialing granitoid (Luo & Miao 2002). Mantle upwelling resulted in crustal extension and lithospheric delamination. Granitic domes and sinistral strike-slip faults were formed. These provided channels and space for the upwelling mantle, such as the NNE –SSW- and NE –SW-trending basic and lamprophyre dykes of 125–110 Ma (Guo et al. 2004), and hydrothermal fluids containing gold. The fluid inclusion measurements described above and isotope data of earlier workers have revealed that ore-forming fluids of Linglong type quartz vein-style and Jiaojia type fault-zone hosted disseminated and stockwork-style gold deposits are consistent throughout the Jiaodong Peninsula, with a similar fluid nature and mineralizing temperature and pressure conditions. The difference between the deposits is demonstrated only by the variety of ore-hosting spaces. As the first-order regional faults formed earlier and were superimposed by multiple deformations, rocks in the fault zones show have higher cataclastic deformation. Mineralization by replacement and infiltration by oreforming fluids took place in the faults and formed disseminated and stockwork-style gold deposits. In the second- or third-order faults, deformation intensities were relatively weak, and rocks in the fault zones have been weakly cracked. Continuous and discontinuous open spaces were formed, which were filled by the ore-forming fluids to form quartz vein-style gold deposits. In the regional-scale faults, the degree of deformation decreased from the central to the outer zone. Disseminated-style gold mineralization occurred in the central fault zone,

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whereas towards the outer zones, fine vein or stockwork-style gold mineralization occurred, followed by quartz –sulphide vein-style gold mineralization. In the margin of Jurassic basins, with deformation divergence of the basement and volcanic rocks, interstratal glide and cracking occurred. Ore-forming fluids ascended along these channels to form the special example of disseminated and stockwork-style (Jiaojia type) mineralization along the interstratal glide zones.

Conclusions (1) Three types of fluid inclusions have been distinguished from quartz and carbonate minerals in gold lodes in the Jiaodong Peninsula: A-type H2O–CO2 inclusions; B-type CO2 –H2O + CH4 inclusions; C-type aqueous H2O inclusions. Fluid inclusion type combinations in each paragenetic sequence are different. (2) Ore-forming fluids depositing gold are consistent throughout the Jiaodong Peninsula, with a similar fluid nature and similar mineralizing temperature and pressure conditions. The fluids are characterized by H2O–CO2 –NaCl + CH4. The optimal mineralizing temperature and pressure are 170– 335 8C and 0.7–2.5 kbar, respectively. (3) Gold and related metallogenic materials were derived from a mixed source, and mantle components were involved in ore-forming processes. The fluids responsible for gold mineralization were derived from fluids degassed from mafic to intermediate magmas. (4) Gold deposits in the Jiaodong Peninsula were formed in the same tectonic–magma– fluid geodynamic background, and are related to the Mesozoic tectonic transition in the eastern North China Craton. Gold metallogeny is only one manifestation of this tectonic inversion. This study was financially supported by the National Science Foundation of China (Grants 40625010 and 40421202). Special thanks are due to the management and staff of Sanshandao, Jiaojia, Linglong, Yingezhuang, Daliuhang, Denggezhuang, Jinqinding, and Penjiakuang Mines for their hospitality during the fieldwork. We appreciate thorough reviews by T. Kusky and Sang Joon Pak, whose insights helped to improve the manuscript.

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A mechanism for transforming buoyant North Chinese cratonic lithosphere to a denser equivalent for delamination J.-F. DENG1, M.-F. ZHOU2, M. F. J. FLOWER3, S.-G. SU1, M.-G. ZHAI4, C. LIU1, G.-C. ZHAO1, X.-G. ZHAO1, S. ZHOU1 & Z.-W. WU1 1

State Key Laboratory of Geological Processes and Mineral Resources, Key Laboratory of Lithosphere Tectonics and Lithoprobing Technology of Ministry of Education, China University of Geosciences, Beijing 10083, China 2

University of Hong Kong, Hong Kong, China

3

University of Illinois at Chicago, Chicago, IL 60607 – 7059, USA (e-mail: [email protected])

4

Institute of Geology and Geophysics, Chinese Academy of Sciences, Beijing 100029, China Abstract: Two models have been proposed to explain lithospheric thinning of North Chinese cratonic lithosphere: (1) thermal erosion or/and chemical metasomatism, causing the lower part of the lithospheric mantle to be transformed into asthenosphere, a mechanism that implies thinning of relatively buoyant lithosphere; (2) the delamination of lithospheric mantle, in whole or part, along with the lowermost crust, as an effect of their increased densities relative to the underlying asthenosphere. This paper explores possible mechanisms whereby buoyant cratonic lithosphere might be transformed into a denser equivalent susceptible to delamination by the convecting asthenosphere. The Yanshan mobile belt in Eastern China developed in response to a combination of subduction and collision. Its apparent ‘counterclockwise’ P– T –t metamorphic evolution suggests that underplated basaltic magma may have heated and, in turn, weakened the cool, rigid crust, allowing for compressional deformation and crustal thickening. Based on three independent lines of evidence (compressional deformation, the record of igneous activity, and lower crustal xenoliths) the thickness of continental crust is estimated to be about 50–65 km. Along with petrological and geochemical studies, thermal modelling shows that large-scale input of asthenospheric basaltic magma leads to granitoid partial melts in the lower crust, and the dominance of highpressure eclogitic products following orogenic thickening may be necessary for eventual delamination to occur.

It is well established that the North China cratonic lithosphere was reactivated during the Jurassic– Cretaceous Yanshanian orogeny, and probably significantly thinned (e.g. Zhang et al. 1983; Liu 1987; Deng 1988; Menzies et al. 1993; Deng et al. 1994, 2004a, b, c). The mechanisms of lithospheric thinning are still not well understood, however, given that buoyant continental cratonic roots are generally believed to be tectonically stable. Two main types of model have been advanced: (1) processes involving thermal erosion or/and chemical metasomatism (e.g. Griffin et al. 1998; Menzies & Xu 1998; Xu 2001; Zhang, H.-F. et al. 2002, 2003; Zhou et al. 2002; Chen et al. 2003; Niu 2005), unrelated to lithosphere density differences, that allow for thinning of the relatively buoyant lithosphere; 2) lithospheric delamination (e.g. Deng et al. 1994, 1996, 2003, 2004a; Gao et al. 1998; Wu et al. 2000, 2002; Zheng et al. 2003; Gao et al. 2004), requiring that the lithosphere is significantly denser than the underlying asthenosphere.

This paper discusses how buoyant cratonic lithosphere in North China may have been transformed into denser equivalent lithologies, prior to its delamination by convecting asthenosphere, as evidenced by Jurassic–Cretaceous magmatic activity and tectonic deformation, along with preliminary thermal modelling results.

Orogenic episodes and structural elements of the Yanshan belt Yanshan mobile belt and regional tectonic setting The ‘Yanshan Mobilization’ was first recognized by Wong (1927), who, on the basis of geological surveys in the areas of Beijing, eastern Hebei and western Liaoning, ascribed the orogeny to Jurassic– Cretaceous time and described the accompanying magmatism, tectonic deformation and mineralization.

From: ZHAI , M.-G., WINDLEY , B. F., KUSKY , T. M. & MENG , Q. R. (eds) Mesozoic Sub-Continental Lithospheric Thinning Under Eastern Asia. Geological Society, London, Special Publications, 280, 317–330. DOI: 10.1144/SP280.17 0305-8719/07/$15 # The Geological Society of London 2007.

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Today, Eastern Asia is made up of a collage comprising the Mongolian, North China and South China Blocks, and intervening Permo-Triassic orogenic belts, accreted during the Mesozoic to the Siberia Craton. During this time, the Yanshan mobile belt was bounded by west-vergent subduction (e.g. Deng et al. 1994, 1996, 2004a, b, c; Van der Voo et al. 1999) and developed in response to the combined effects of subduction and colliding terranes, as reflected by its deformational style and calc-alkaline magmatic affinity (see Figs 1 and 2). Several key aspects remain problematic, however. These include, for example: (1) the extent to which relatively cold, rigid cratonic crust was able to undergo compressional deformation and thickening; (2) the significance of coeval extensional basins (e.g. Griffin et al. 1998; Menzies & Xu 1998; Chen et al. 2002, 2005; Ge et al. 2002); (3) the extent to which crustal thickening resulted from tectonic processes (e.g. Deng et al. 1996; Davis 2003; Zhang et al. 2003) as opposed to the effects of magmatic underplating (e.g. Xu et al. 2004). Moreover, there are two contrasting interpretations as to when compressional tectonics gave way to extension; the first prefers the Late Jurassic (J3) (e.g. Lu et al. 1997; Wang & Zhang 2001), and the second, the early Cretaceous (K1) (e.g. Davis 2003; Deng et al. 2003).

Orogenic episodes of the Yanshan belt In Table 1 we present a schematic summary of the main magmatic–tectonic events associated with the Jurassic – Cretaceous Yanshanian orogenic belt (Deng et al. 2004a, and references therein). Five orogenic episodes are recognized on the basis of unconformities: (1) a precursive, or pre-orogenic (J1) episode; (2) an early (J2) orogenic episode; (3) a peak (J3) orogenic episode; (4) a late (K11, early early Cretaceous) orogenic episode; (5) a postorogenic (K12, late early Cretaceous) episode. Except for the last of these, each corresponds to a discrete magmatic–tectonic cycle progressing from early volcanism to sedimentation, plutonic activity, compressional regional metamorphism, and, finally, uplift and erosion. Each cycle appears to record a counterclockwise metamorphic P–T –t path, indicating that initial mantle-derived heat transfer precedes crustal thickening, uplift and erosion, to be followed by crustal decompression. According to this scheme, precursive, magmatic underplating would serve to weaken the preexisting cratonic crust, producing rheological conditions conducive to compressional deformation and crustal thickening; the metamorphic record indicates progressive increases in temperature and pressure from J1 to J3 (Table 1), accompanied

(presumably) by crustal thickening, prior to the orogenic peak in the late Jurassic (J3).

Structural elements of the Yanshan orogenic belt Whereas the main geological events are marked by tectonic, magmatic, metamorphic and sedimentary processes, the principal lithological structural elements comprise folding, faulting and schistosity (Deng et al. 2004b, and references therein). These are shown in Table 1 in relation to each orogenic episode, indicating their genetic relationships, as follows. (1) Pre-orogenic extension (J11) is characterized by mafic volcanic rocks and coal-bearing sedimentary formations, initial orogenic compression (J12) being associated with molasse-like formations and relatively high-pressure–lowtemperature metamorphism. (2) Both early (J2) and peak (J3) orogenic episodes are characterized by compressional deformation along with felsic and silicic igneous activity, molasse formation, and, at deeper levels, amphibolite-facies metamorphism. (3) The late orogenic episode (K11) is associated with local compressional deformation and dominantly mafic igneous activity, locally including alkali syenite and lacustrine formations. (4) Post-orogenic (K12) features, in contrast, include normal faulting, coal-bearing formations, bimodal dyke swarms, alkali granites, and metamorphic core complexes. From the above, it is clear that structural element associations differ from one episode to the next, reflecting the overall thermal and tectonic evolution of the Yanshan belt. In other words, as crustal heating progresses, compressional deformation peaks (J3), giving way thereafter, following a transitional stage (K11), to extensional and finally transtensional deformation (K12), at which point progressive cooling and crustal uplift occur.

Multi-directional tectonic vergence of the Yanshan orogenic belt Referring to it as the Yanshan fold and thrust belt, Davis et al. (2001) and Davis (2003) drew attention to the temporal and geometric complexity of Mesozoic contraction, especially with regard to the coexistence of north- and south-directed thrust faulting and the lack of a consistent sense of tectonic vergence such as typifies most foreland fold–thrust belts. Our present studies indicate a succession of five compressional deformational events, of contrasting vergence, in the Yanshan belt (Table 1). During both late J1 and late J2 time, east–west-striking compressional thrusting was north-directed,

Fig. 1. Simplified map showing the distribution of volcanic rocks in the Yanshan belt. The inset map shows the research area.

TRANSFORMING BUOYANT LITHOSPHERE TO DENSE 319

Fig. 2. Simplified map showing the distribution of intrusive rocks in the Yanshan belt. 1, K2 intrusion; 2, K1 intrusion; 3, J3 intrusion; 4, J2 intrusion; 5, J1 intrusion; 6, granite; 7, alkaline granite; 8, gabbro; 9, monzonite; 10, monzodiorite; 11, diorite; 12, syenite; 13, alkaline syenite; 14, normal fault; 15, reversed fault; 16, geological boundary; 17, intrusion name; 18, intrusion age.

320 J.-F. DENG ET AL.

T3

J1

J2

Nandaling

Jiulongshan Xiahuayuan

Houcheng Tiaojishan

Molasse-like Coal-bearing formation (tuff layers, R, 180 – 178 Ma) B – TB – BTA –TA – T, 196– 184 Ma

Molasse TA – T– D –R, 173 – 161 Ma

h – g(?)199 – 196 Ma

h– Qh – g, 174– 151 Ma

y– hd– h– Qh –g 148 – 136 Ma

y –hy– hd –h – A/j –j – g 133 – 127 Ma

A/g – g (118– 119 Ma), bimodal dyke swarm(120 – 114 Ma)

Intrusive rock assemblage

Indosinian south-vergent thrust faulting (pre-180 Ma or pre-199 Ma), Pingquan– Chengde – Chicheng mylonite (c. 211 Ma)

Late J1 ENE – WSW-trending folding and thrusting (Western Hills of Beijing), and chloritoid– staurolite –kyanite metamorphism (,178 – 180, 175 Ma?)

Late J2 – early J3 northward thrust (161 – 148 Ma) (Chengde, Shisanling, Xinglong) and amphibolite-facies metamorphism(159 – 151 Ma)

Late J3 NE – SW-trending thrusting (138 –136 Ma) J3 southward thrusting (144 – 140 Ma) (Sihetang, Gubeikou) and amphibolite-facies metamorphism(144 – 138 Ma)

Late K12 metamorphic core complex, Yunmengshan, 119 – 114 Ma, and normal faulting Early K11 localized reverse faulting with NW – SE trend (,125 Ma)

Tectonic deformation phase and metamorphism

After Deng et al. (2004a, and references therein), with supplements. B, basalt; TB, trachybasalt; BTA, basaltic trachyandesite; TA, trachyandesite; D, dacite; T, trachyte; R, rhyolite; y, gabbro; hy, monzogabbro; hd, monozodiorite; h, monzonite; Qh, quartz monzonite; j, syenite; A/j, alkaline syenite; g, granite; A/g, alkaline granite; zig zag line indicates angular unconformity. The classification and nomenclature of volcanic and intrusive rocks are after Le Maitre (1989) and Middlemost (1994), respectively. The stage ages are after Remane et al. (2000).

203 Ma

175 Ma

154 Ma

J3

Molasse TA – T– D –R, 148 – 140 Ma

B – TB –BTA, 135 – 121 Ma

Yixian

Shouwangfen Zhangjiakou

Lake-basin formation

Jiufutang

135 Ma

Coal-bearing formation

Fuxin

K1

Molasse-like(?)

Sunjiawan

K2

96 Ma

Volcanic and sedimentary formation

Stratigraphy

Time

Table 1. A preliminary scheme of the Yanshanian (Jurassic – Cretaceous) magmatic– tectonic event sequence of the Yanshan orogenic belt, North China

TRANSFORMING BUOYANT LITHOSPHERE TO DENSE 321

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J.-F. DENG ET AL.

whereas J3 (144 –140 Ma) thrusting, also east– west striking, is south-vergent. By late J3 time (138 – 136 Ma) the strike of compressional thrusting had changed abruptly to NE–SW, with both NW- and SE-directed vergence. Moreover, late Early Cretaceous (K11) thrust faulting reflects a further abrupt change in strike to NW –SE, showing NE- and SW-directed vergence. It is noteworthy that the direction of both the long axis of single intrusive bodies and the extension of syntectonic intrusive rock assemblages are consistent with the trend of coeval compressional thrusting (Deng et al. 2005), suggesting a genetic feedback relationship between magmatism and compressional deformation. We can further consider that the observed variations in tectonic vergence are genetically related to the configuration of circum-East Asian subduction systems and, in turn, their accretion to the Siberian Craton. However, this relationship needs further detailed study.

Independent lines of evidence for compressional thickening of Yanshanian crust: igneous petrology, geochemical character, and lower crustal xenoliths Compressional deformation The evidence for multi-episodic, multi-directional compressional deformation discussed above (Table 1) is believed to be one of the most important lines of evidence for crustal thickening. However, estimation of the approximate extent of crustal thickening cannot be based on the observed compressional deformation alone but requires input from petrological and geochemical studies.

High-pressure trachyte and syenite, including trachyandesite, monzonite and quartz monzonite Considering petrological phase equilibria, Wyllie (1977, 1984) concluded that trachytic and syenitic magmas rather than granite are formed by partial melting at the base (.50 km) of thickened continental crust. On the basis of available experimental data, Deng et al. (1998, and references therein) were able to define the water-undersaturated liquidus surface in P–T space for trachyte and syenite (Fig. 3), inferring that these magmas equilibrate with eclogitic residua at pressures of 15 kbar or more. This observation is taken to account for the lack of negative Eu anomalies in high-pressure trachyte and syenite melts. On the other hand, adakitic magmas, generated as partial melts of subducting

Fig. 3. Schematic diagram of water-undersaturated liquidus surface for trachyte and syenite showing liquidus and near-liquidus minerals ( after Deng et al. 1998, and references therein). Pl, plagioclase; Qz, quartz; Ct, coesite; Cpx, clinopyroxene; Jd, jadeite and jadeitic pyroxene; Hb, hornblende; Ga, garnet.

oceanic slabs, are also assumed to have formed at pressures 15 kbar and, although they show relatively high Sr/Y ratios (e.g. Drummond et al. 1996), they are also characterized by the lack of negative Eu anomalies This raises the question of whether trachytes and syenites in the Yanshanian belt formed at the base of thickened continental crust, or as (adakitic) slab melts. From Table 2 and Figure 4, it may be seen that: (1) with the exception of J1, igneous rocks from J2 to K11 (e.g. J2 (WM03-10, TA), J2/J3 (Wy-26, Qh), J3 (Wy-60, Qh), J3 (XS-11, T), K11(Wy-4, j) show high-pressure trachytic and syenitic affinity; (2) these high-pressure Yanshanian magmatic products differ from Cenozoic adakites in terms of their respective total alkalis–silica (TAS) classification and nomenclature, their covariance of SiO2 –MgO, SiO2 – FeO/MgO and SiO2 –K2O relations, Ni contents, and K2O/Na2O ratios; (3) the higher MgO and Ni contents and lower values of FeO/MgO, K2O and K2O/Na2O observed in adakites result from both the interaction between slab melts and mantle wedge peridotite and the midocean ridge basalt (MORB)-like source of the former (e.g. Drummond et al. 1996); (4) in contrast to adakites, the lower MgO and Ni contents, and higher values of FeO/MgO, K2O and K2O/Na2O for the high-pressure Yanshanian trachytes and syenites are consistent with their formation at the base of thickened continental crust, with no history of interaction with mantle peridotite. As is well known, values of K60 (K2O wt%, at SiO2 60%) may be used to establish a rough estimate of crustal thickness (e.g. Condie 1982).

57.23 (TA) 2.24 3.04 (Th) HKCA

0.39 838 16.5 50.79 0.84 15.1

62.85 (T) 1.35 3.66 (Th) HKCA

0.42 336 15.9 21.13 0.89 2.63

0.94 606 17.1 35.44 0.76 1.56

64.67 (T) 0.91 4.32 (Th) SH 0.91 238 18.9 12.59 0.78 6.01

66.16 (T) 0.95 4.07 (Th) SH 0.51 794 23.8 33.36 0.83 2.09

59.8 (TA) 1.78 3.31 (Th) HKCA

WM03-14 WM03-10 WM03-21 WM03-3 WM03-17 J3 J1 J2 J2 J3* 145 Ma 184 Ma 161 Ma 167 Ma 137 Ma

0.28 1004 21.7 46.27 1.0 2.59

63.42 (T) 0.99 5.21 (Th) MKCA

XS-11 J3 138 Ma

0.89 690 40.5 17.04 0.77 15.7

59.62 (h) 2.39 2.48 (CA) SH

WY-62 J1 197 Ma

1.24 370 19.4 19.07 0.56 11.8

67.98 (Qh) 1.37 2.69 (CA) HKCA

YS-23 J2 159 Ma

0.71 899 16.7 53.83 0.88 25.61

61.12 (Qh) 2.88 1.91 (CA) HKCA

WY-26 J2/J3 153 Ma

1.13 842 8.01 105.12 1.09 4.77

68.69 (Qh) 1.13 2.54 (CA) HKCA

WY-60 J3 138 Ma

Intrusive rocks

1.07 467 18.9 24.71 0.57 4.20

69.86 (Qh) 0.80 3.59 (CA) HKCA

WM03-26 J3 136 Ma

0.82 1348 15.4 87.53 1.99 2.72

59.85 ( j) 1.45 2.85 (CA) SH

WY-4 K11 124 Ma

39.0

0.35 869 9.5 91.47

63.84 (D) 2.47 1.70 (CA) MKCA

Average Cenozoic adakite2 (n ¼ 140)

Source: 1This paper, unpublished data.2Drummond et al. (1996). The age is SHRIMP zircon U –Pb age of the sample, however, for example X5-11 the age is that of the same strata nearby. The classification and nomenclature of volcanic and intrusive rocks are after Le Maitre (1989) and Middlemost (1994), respectively. T, trachyte; TA, trachyandesite; D, dacite; h, monzonite; Qh, quartz monzonite; j, syenite; Th, CA are tholeiitic and calc-alkali series after SiO2 –FeO/MgO relation from Miyashiro (1974). The FeO/MgO, K2O/Na2O and Sr/Y are calculated in this paper from the average composition of the Cenozoic adakite after Drummond et al. (1996).

SiO2 – K2O relation K2O/Na2O Sr (ppm) Y (ppm) Sr/Y dEu Ni (ppm)

MgO (wt%) FeO/MgO

SiO2 (wt%)

No.: Age:

Volcanic rocks

Yanshan belt1

Table 2. Some chemical parameters for the igneous rocks with 57 and ,70 wt% SiO2 from the Yanshan belt and the average Cenozoic adakite

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J.-F. DENG ET AL.

Fig. 4. MORB-normalized trace-element patterns of the Yanshanian rocks (after Pearce 1983).

Accordingly, igneous products in the Yanshan belt being dominated by high-K calc-alkaline (HKCA) affinity (Deng et al. 1996) may be taken to indicate crustal thickening up to about 50–65 km (Deng et al. 1996, 2003), consistent with observed values of Sr/Y ratio and dEu (Table 2, Fig. 2c).

Lower crustal xenoliths Recently, Fan et al. (1998) provided the first zircon U– Pb dating of 140– 120 Ma for mafic lower

crustal xenoliths in the Cenozoic Hannuoba basalts, and Wilde et al. (2003) provided zircon U –Pb age dates indicating a peak age between 180 Ma and 80 Ma for mafic and felsic granulite and pyroxenite xenoliths. Chen et al. (2001) presented thermobarometric data for nine-samples of Cr-spinel garnet pyroxenite (+ olivine), indicating that equilibration conditions ranged from 15 kbar and 993 8C to 25.3 kbar and 1265 8C. Most pressure estimates were between 16 and 22 kbar, approximately equivalent to depths of 50– 65 km; these

TRANSFORMING BUOYANT LITHOSPHERE TO DENSE

values are generally consistent with estimates of crustal thickening derived from K60 values in the Yanshanian HKCA igneous series (Deng et al. 1996, 2003).

Subduction-related tectonic setting: evidence from the igneous rocks Pitcher (1993) has shown that oceanward and inland sides of active continental cordilleras are characterized respectively by tonalite –granodiorite and granodiorite–granite associations. Brown (1982) also showed that, according to the alkali –lime index of Peacock (1931), batholith granitoids of western North American and the New Guinea mobile belt are calc-alkaline (CA), whereas the New Guinea arc margins are of alkali –calcic (AC) to alkaline affinity. Based on the absence of a tonalite – granodiorite association and the dominance of quartz monzonite– granite, alkali –calcic (AC) and HKCA associations, the Yanshan belt is believed to have been located on the inland side of the Mesozoic western Pacific margin (e.g. Wang & Mo 1995; Deng et al. 1996; Davis 2003). Except for postorogenic (K12) lithologies, this interpretation is supported by the Yanshanian igneous trace element distributions (Fig. 4). The recent recognition of J3 and K11 boninitic (or high-magnesium andesite) magmatic products and high-Mg adakite in western Liaoning Province (Zhang, Q. et al. 2003; Gao et al. 2004; Zhang et al. 2005) provides further support for the hypothesis that the Yanshan belt represents a Mesozoic continental arc –forearc complex. Their strong enrichment in large ion lithophile elements (LILE) relative to high field strength elements (HFSE) (Fig. 4) confirms that Yanshanian magmatic sources were contaminated by subduction-related aqueous fluids and/or melts, consistent with the regional tectonic setting proposed. Although the igneous activity is clearly typical of active continental margin settings, its westernmost occurrence is more than 1000 km from the presumed location of the oceanic trench, implying that Mesozoic subduction effects were unusually far-reaching. However, despite significant crustal shortening and thickening produced by multiepisodic and multi-directional compressional deformation, this effect can be attributed to subsequent extensional deformation and vertical thinning that has occurred since the late Early Cretaceous (K21). Taking the Sea of Japan, the Bohai and Yellow Seas, and rift-related basins in Northern China into account, the effective distance between the continental arc and deep-sea trench would be substantially reduced. Seismic tomographic studies

325

(e.g. Fukao et al. 1992; Van der Voo et al. 1999) indicate the presence of subducted slab remnants at the mantle transition zone, providing additional support for Mesozoic subduction beneath the North China Block.

Basalt magmatic underplating and crust –mantle interaction in the Yanshan belt Generation of mafic magma Experimental studies have established that basaltic and boninitic (high-Mg andesite) partial melts reflect respectively (1) anhydrous or nearanhydrous partial melting of ‘fertile’ lherzolite (LH) sources and (2) hydrous partial melting of ‘refractory’ harzburgite (or Cpx-harzburgite) (HZ) sources, under more or less similar pressure conditions (e.g. Kushiro 1990; Hirose & Kushiro 1993; Hirose 1997; Falloon & Danyushevsky 2000). Table 3 summarizes the SiO2 –MgO covariation of Yanshanian gabbros and high-Mg andesites associated with the North China Block along with that of experimental partial melts of variably fertile LH and HZ source peridotites. Compared with partial melts of LH, those of HZ have significantly higher MgO for equivalent SiO2 content. Given this relationship, we infer that the Yanshanian mafic magmas, of both gabbroic and high-Mg andesite composition, are produced by partial melting of lherzolite rather than harzburgite sources. As discussed above, the North China cratonic mantle lithosphere is mainly composed of harzburgite, implying that the Yanshanian mafic magmas tap asthenospheric rather than lithospheric mantle sources. Experimental data bearing on high-Mg andesite magmas appear to indicate hydrous rather than anhydrous sources, suggesting the presence of slab-derived H2O in a supra-subduction mantle wedge, consistent with the inferred regional tectonic setting of the circum-East Asia region.

Two types of silicic (SiO2 .70 wt%) magmas: hybrid and non-hybrid granites Experimental studies of hybrid compositions comprising 50% high-Al olivine tholeiite (HAOT) and 50% biotite gneiss (BG), at 1000 8C and pressures of 0.5, 0.7, 1.0, 1.2 and 1.5 GPa (Patin˜o Douce 1995) yielded between 32 and 38 wt% of H2Oundersaturated granitic melts (SiO2 .70 wt%) and allow for a plausible crustal assimilation model. In Table 4, we can observe three features. A nonhybrid silicic melt derived from biotite gneiss (BG) (Patin˜o Douce & Beard 1995) has essentially

326

J.-F. DENG ET AL.

Table 3. Comparison between the Yanshanian (J–K) mafic igneous rocks of North China and the experimental partial melts from peridotites Rock or melt Pingshun & Jinan, y Qipanyan, Beijing y LH-melt HZ-melt HMA western Liaoning LH-melt HZ-melt

SiO2 (wt%)

MgO (wt%)

Source

49 48 48–49 51 55 55 54

16 17 10 –16 21 8 6 –11 .22

Deng et al. (2003) This paper Hirose & Kushiro (1993) Falloon & Danyushevsky (2000) Zhang et al. (2003) Kushiro (1990); Hirose (1997) Falloon & Danyushevsky (2000)

y, gabbro; HMA, high-Mg andesite; LH-melt, melts formed from lherzolite; HZ-melt, melts formed from harzburgite.

the same SiO2 content (.70 wt%) as that of the hybrid source (BG þ HAOT), although the latter is more CA in character (after Yogodzinski et al. 1995; Frost et al. 2001) (i.e. shows higher MgO and CaO, and lower FeO/MgO) as compared with the non-hybrid silicic melt. In accord with this distinction, many of the high-SiO2 (.70 wt%) Yanshanian rhyolites and granites are believed to be silicic melts of hybrid sources. Thus, the Yanshanian magmatic products (Table 1) are believed to have formed from basalt magmatic underplating and crust–mantle interaction, as well as magma mixing and fractional crystallization (Deng et al. 1996).

Thermal modelling As is well known, large-scale production of continental granitoid magmas reflects underplating of the lower crust by mantle-derived basaltic magma. Our preliminary thermal modelling, following the method of Bergantz (1989), shows that ratios of (1) total granitic melt and (2) total extractable granitic magma, generated in tonalitic lower crust to the total amount of underplated basalt, are about 0.12 and 0.03, respectively (Liu 2004). These values imply that large volumes of basaltic magma are required to produce silicic magmas in cold tonalitic craton lower crust. For example, production of a 1 km thick granitic body would require about 33 km of underplating basalt to crystallize. In turn, eclogitization of such volumes of basalt in the vicinity of the Moho would increase the ambient density of the lower lithosphere sufficiently to allow for delamination.

Lithosphere delamination Conditions necessary for delamination Kay & Kay (1993) showed that for regions with relatively thin crust (,50 km), lower crust of any composition has a lower density than that of the mantle, in contrast to regions of compression-related

crustal thickening, where lower crustal basalts undergo large density increases following their transition to eclogite. Thus, two conditions, both satisfied between J2 and K11 in the Yanshanian belt, are necessary to allow for delamination: (1) basalt magmatic underplating; (2) compressional deformation resulting in crustal thickening up to .50 km.

Shallow crustal response to lithosphere delamination The process of delamination is favoured because both the rate of conduction and radioactive heating effects would be insufficient to affect the lithospheric mantle (i.e. convert it to asthenosphere) on the scale required over short time intervals corresponding to observed tectonic transitions in orogenic zones. Delamination is viewed as a rapid, ‘catastrophic’ event, in contrast to ‘thermal thinning’, which, (with no advection of heat) would probably be unrealistic on a short time scale (Kay & Kay1993). It is clear from the Yanshanian magmatic–tectonic record (Table 1) that the observed sequence of compressional deformation, uplift and erosion, and volcanism occurred over relatively short time intervals, consistent with the notion of ‘catastrophic’ delamination rather than a protracted period of thermal erosion. Kay & Kay (1993) also observed that total crustal extension is relatively limited when associated with upwelling asthenosphere, a factor that is often downplayed or ignored by models proposing crustal extension as a direct result of delamination. The possibility of multiple delamination events during the (J2 – K11) Yanshanian (Table 1) is consistent with the Kay & Kay (1993) model, given that the region was unaffected by extensional deformation until the post-orogenic episode (K12).

Lithospheric delamination model From the above discussion, we conclude that a lithosphere delamination model is able to explain

64.6 4.7 2.1 1.66 3.5

46.9 11.0 11.9 0.75 29.6

HAOT

1

1

55.7 7.8 7.0 1.03 23.0

BG þ HAOT

Starting materials

72.03 (70.2– 74.09) 0.93 (0.57– 1.07) 2.33 (1.62– 3.23) 1.87 (1.26– 2.33) 5.52 (4.28– 6.89)

2

73.62 (70.20 – 76.40) 0.47 (0.30 – 0.84) 1.23 (0.64 – 2.18) 3.64 (2.01 – 6.52) 6.69 (3.30 – 9.29)

Non-hybrid (n ¼ 24)

Experimental melts Hybrid (n ¼ 6)

1

73.10 (70.11– 76.93) 0.57 (0.24– 2.98) 0.92 (0.07– 6.22) 3.87 (0.32– 11.27) 7.83 (0.37– 9.94)

74.48 (71.77 –76.46) 0.07 (0.01 – 0.19) 0.41 (0.02 – 0.65) 22.58 (8.19 – 98.42) 8.83 (8.00 – 9.51)

Non-hybrid4 (n ¼ 9)

Yanshan belt Hybrid (n ¼ 39)

3

Data sources: 1Patin˜o Douce (1995); 2Patin˜o Douce & Beard (1995); 3This paper; Bai et al. (1991); Bao et al. (1995); 4This paper, Table 3. BG, biotite gneiss; HAOT, high-Al olivine tholeiitic glass; BG þ HAOT, 50% BG þ 50% HAOT, using hybrid experiment as bulk assimilation model. The experimental hybrid and non-hybrid silicic melts are formed from the partial melting of the BG þ HAOT and the BG starting materials, respectively; values in parentheses are the compositional range. The FeO/MgO ratio, (Alk –CaO) and average compositions are calculated in this paper.

SiO2 (wt%) MgO (wt%) CaO (wt%) FeO/MgO Alk –CaO (wt%)

BG

1

Table 4. Some chemical parameters for the high-SiO2 (.70 wt%) melts from both the non-hybrid and the hybrid experiments, and the high-SiO2 (.70 wt%) silicic rhyolite and granite rocks of the Yanshan belt

TRANSFORMING BUOYANT LITHOSPHERE TO DENSE 327

328

J.-F. DENG ET AL. We would like to thank the reviewers J. Encarnacion and T. Kusky for substantial comments on the manuscript. This work was supported by the NSF of China (Award 40234048), the Chinese Geological Survey (Award 20001010202), and the Ministry of Science and Technology (Award 2001cb711002).

References

Fig. 5. Delamination model showing the evolution of the lithosphere beneath North China during Jurassic– Cretaceous time (after Deng et al. 2003, with supplements). L, cratonic lithosphere; A, asthenosphere; B, basaltic magma; YSLN, Yanshan–western Liaoning area; L2, lithosphere delamination; DXAL, Daxinganling; SP, Siping; YB, Yanbian; L3 new lithosphere formed from cooling of the asthenosphere Pre-Paleogene.

the major stages of the evolving Yanshanian orogenic belt (Fig. 5), as follows. (1) Pre-orogenic phase (J1) (Fig. 5a): having undergone extensional deformation, the cratonic lithosphere ruptures along previous, MesoProterozoic, rift zones. This allows uprise of basaltic magma from the asthenosphere, which leads to underplating at the base of the crust and volcanism in the Yanshan–western Liaoning area. (2) Syn-orogenic phase (J12 initial-orogenic to K11 late-orogenic episodes), lithosphere delamination (Fig. 5b): multi-episodic contractional deformation results in thickening of crust and underplated basalt. Eclogitization of the latter leads to increased density and, in turn, catastrophic lithosphere delamination. Eruptive and intrusive active becomes widespread to the east and NE of the Ordos block, including Daxinganling, the Yanshan –western Liaoning area, Siping and Yanbian. (3) Post-orogenic phase (K12) to Pre-Paleogene (Fig. 5c): formation of composite lithosphere.

B AI , Z. M., X U , S. Z. & G E , S. W. 1991. The Badaling Granitic Complex. Geological Publishing House, Beijing [in Chinese]. B AO , Y. G., B AI , Z. M., G E , S. W. & L IU , C. 1995. Yanshanian Volcanic Geology and Rocks. Geological Publishing House, Beijing [in Chinese]. B ERGANTZ , G. W. 1989. Underplating and partial melting: implications for melt generation and extraction. Science, 245, 1093–1095. B ROWN , G. C. 1982. Calc-alkaline intrusive rocks: their diversity, evolution, and relation to volcanic arcs. In: T HORPE , R. S. (ed.) Andesites, Wiley, Chichester, 437–461. C HEN , B., Z HAI , M. G. & S HAO , J. A. 2002. Petrogenesis and significance of the Mesozoic North Taihang complex: major and trace element evidence. Science in China, Series D, 32, 896– 907 [in Chinese]. C HEN , B., J AHN , B. M. & Z HAI , M. G. 2003. Sr–Nd isotopic characteristics of the Mesozoic magmatism in the Taihang– Yanshan orogen, North China Craton, and implications for Archean lithosphere thinning. Journal of the Geological Society, London, 160, 963–970. C HEN , B., T IAN , W., Z HAI , M. G. & A RAKAWA , Y. 2005. Zircon U –Pb geochronology of the Mesozoic magmatism in the Taihang mountain and other places of the North China Craton, with implications for petrogenesis and geodynamic setting. Acta Petrologica Sinica, 21(1), 13–24 [in Chinese with English abstract]. C HEN , S. H., O’R EILLY , S. Y., Z HOU , X. H., ET AL . 2001. Thermal and petrological structure of the lithosphere beneath Hannuoba, Sino-Koreau Craton, China: evidence from xenoliths. Lithos, 56, 267– 301. C ONDIE , K. C. 1982. Plate Tectonics and Crustal Evolution. Pergamon, New York. D AVIS , G. A. 2003. The Yanshan belt of North China: tectonics, adakitic magmatism, and crustal evolution. Earth Science Frontiers, 10(4), 373– 384. D AVIS , G. S., Z HANG , Y. D., W ANG , C. ET AL . 2001. Mesozoic tectonic evolution of the Yanshan fold and thrust belt, with emphasis on Hebei and Liaoning Provinces northern China. In: H ENDRIX , M. S. & D AVIS , G. A. (eds) Paleozoic and Mesozoic Tectonic Evolution of Central Asia: from Continental Assembly to Intracontinental Deformation.Geological Society of America. Geological Society of America, Memoirs, 194, 171– 197. D EFANT , M. J., X U , J. F., K EPEZHINSKAS , P. ET AL . 2002. Adakites: some variations on a theme, Acta Petrologica Sinica, 18(2), 129–142. D ENG , J. F. 1988. Continental rift magmatism and deep process. In: C HI , J.-S. (ed.) Cenozoic Basalt and Upper Mantle of Eastern China. China University of Geosciences Press, Wuhan, 201– 218 [in Chinese].

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W YLLIE , P. J. 1984. Constraints imposed by experimental petrology on possible and impossible magma sources and products. Philosophical Transactions of the Royal Society of London, Series A, 310, 439– 456. X U , Y. G. 2001. Thermal–tectonic destruction of the Archean lithospheric keel beneath the Sino-Korean Craton in China: evidence, timing and mechanism. Physics and Chemistry of the Earth (A), 26, 747–757. X U , Y. G., H UANG , X. L., M A , J. L. ET AL . 2004. Crust– mantle interaction during the tectonic–thermal reactivation of the North China Craton: constraints from SHRIMP zircon U–Pb chronology and geochemistry of Mesozoic plutons from western Shandong. Contribution to Mineralogy and Petrology, 147, 750–767. Y OGODZINSKI , G. M., K AY , R. W. & V OLYNETS , O. N. 1995. Magnesian andesite in the western Aleutian Komandorsky region: implications for slab melting and processes in the mantle wedge. Geological Society of America Bulletin, 107(5), 505–519. Z HANG , H., L IU , X.-M., L I , Z.-T. ET AL . 2005. Early Cretaceous large-scale crustal thinning in the Fuxin– Yixian basin and adjacent area in western Liaoning. Geological Review, 51(4), 360– 372 [in Chinese with English abstract]. Z HANG , H.-F., S UN , M. & Z HOU , X.-H. 2002. Mesozoic lithosphere destruction beneath the North China Craton: evidence from major-, trace-element and Sr– Nd– Pb isotopic studies of FangCheng basalts. Contributions to Mineralogy and Petrology, 144, 241–253. Z HANG , H. F., S UN , M. & Z HOU , X. H. 2003. Secular evolution of the lithosphere beneath the eastern North China Craton: evidence from Mesozoic basalts and high-Mg andesites. Geochimica et Cosmochimica Acta, 67(22), 4373–4387. Z HANG , Q., W ANG , Y. & L IU , H.-T. 2003. On the space– time distribution and geodynamic environments of adakites in China, Annex: controversies over differing opinions for adakites in China, Earth Science Frontiers, 10(4), 385–400 [in Chinese with English abstract]. Z HANG , W. Y., Z HANG , K., Z HAO , Y. G. ET AL . 1983. The Mesozoic and Cenozoic geotectonic characteristics and dynamical model of the lithosphere in North China faultblock region. Acta Geologica Sinica, 57(1), 33–42 [in Chinese with English abstract]. Z HENG , J. P., S UN , M., L U , F. X. & P EARSON , N. 2003. Mesozoic lower crust xenoliths and their significance in lithospheric evolution beneath the Sino-Korean Craton. Tectonophysics, 361, 37–60. Z HOU , X.-H., S UN , M. & Z HANG , G.-H. 2002. Continental crust and lithospheric mantle interaction beneath North China: isotopic evidence from granulite xenoliths in Hannuoba, Sino-Korean craton. Lithos, 62, 111–124.

Lithospheric thinning in eastern Asia; constraints, evolution, and tests of models T. M. KUSKY1, B. F. WINDLEY2 & M.-G. ZHAI3 1

Department of Earth and Atmospheric Sciences, St. Louis University, St. Louis, MO 63103, USA (e-mail: [email protected]) 2

Department of Geology, University of Leicester, Leicester LE1 7RH, UK

3

Key Laboratory of Mineral Resources, Institute of Geology and Geophysics, Chinese Academy of Sciences, Beijing 100029, China Abstract: The North China Craton (NCC) is the only place currently recognized where an Archaean craton developed a continental root in the Archaean, and subsequently lost half of that root in younger tectonism. In this volume, various authors have advanced different models of root loss, and provided geological, geophysical and geochemical data that help constrain the geometry and timing of root loss. Understanding why and how roots are lost may help us understand how often this process may have occurred in the geological past, and how much lithospheric material has been recycled to the convecting mantle through this mechanism, potentially drastically changing our current understanding of crustal growth rates and processes. With current data, there are several equally plausible possibilities that require further data collection for testing. There are several possible tectonic triggers that may have caused half the root to be lost, acting either separately or together. These include collisional, extensional, plume-related, fluid-weakening, spontaneous, and more complex hybrid mechanisms. We also do not know why only the eastern half of the root was lost, and not the root from beneath the whole craton. One tantalizing idea is that the root grew independently, by tectonic underplating of subducted buoyant oceanic lithosphere, beneath the previously separate eastern and western halves of the craton by 2.5 Ga, with modification at 1.8 Ga. If so, perhaps only the eastern half of the root was lost in younger tectonism because there was some physical or geometric difference between the two halves. Alternatively, collisional or subduction-related tectonic processes acting only on the Eastern Block may have caused the disruption of the tectosphere there in the Mesozoic. The timing of and mechanism for loss of the root is not uniquely resolvable with current data, but a solution to the problem is in reach. Possible triggering mechanisms include, but are not limited to, collision of the South China (Yangtze) and North China Cratons in the Triassic, the India– Asia collision, closure of the Solonker and Mongol – Okhotsk oceans, Mesozoic subduction of the Pacific plate beneath Eastern China, impingement of mantle plumes, mantle hydration from long-term subduction, and several rifting events. In this concluding review, we link studies of crustal tectonics with investigations aimed at determining the nature and timing of the formation and loss of the root, to better understand mechanisms of continental root formation, evolution and recycling –removal.

Cratons are structurally complex regions that attained prolonged stability (1 Ga) within the continents and thus, by definition, are Precambrian in age (Fig. 1). Indeed, most formed in the Archaean (Rudnick 1995; Windley 1995; Kusky & Polat 1999). Some first-order observations about Archaean cratons include the following. (1) Many are preserved in the centre of continental masses, surrounded by younger fold belts (Fig. 1). (2) Surface heat flow is typically low within cratons, averaging 41 + 11 mW m22 (Morgan 1984; Nyblade & Pollack 1993; Artemieva & Mooney 2001).

(3) Cratons are underlain by regions of fast P and S velocities, extending to depths of at least 200–250 km, and perhaps deeper (Jordan 1975; James et al. 2001; Gung et al. 2003). (4) Kimberlites erupted within Archaean cratons carry diamonds interpreted to be derived from the lithospheric mantle, consistent with the presence of a cool and thick keel (Boyd & Gurney 1986). (5) Mantle xenoliths from craton-penetrating kimberlites are highly refractory peridotites, which experienced large degrees of melt extraction (Boyd 1989), followed by variable degrees of metasomatic overprinting (Erlank et al. 1987).

From: ZHAI , M.-G., WINDLEY , B. F., KUSKY , T. M. & MENG , Q. R. (eds) Mesozoic Sub-Continental Lithospheric Thinning Under Eastern Asia. Geological Society, London, Special Publications, 280, 331–343. DOI: 10.1144/SP280.18 0305-8719/07/$15 # The Geological Society of London 2007.

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Fig. 1. Map showing the worldwide distribution of Archaean cratons, and the location of the North China Craton (after Kusky & Polat 1999).

(6) Osmium isotope investigations of cratonic peridotites demonstrate that the lithospheric keel formed more or less coincident with the overlying crust (at least within the precision of Os models ages, e.g. +200 Ma) (Pearson et al. 1995a, b; Chesley et al. 1999; Hanghoj et al. 2001; Irvine et al. 2003). Thus, in general, Archaean cratons are characterized by cold, thick, structurally complex lithosphere whose density is offset by its refractory composition, giving rise to chemical buoyancy (Jordan’s tectosphere hypothesis (Jordan 1975, 1981)). The lack of high elevation or free-air gravity anomalies over cratons suggests that they are in isostatic equilibrium and have a density profile that matches those of off-craton regions (i.e. Jordan’s isopycnic hypothesis is generally correct). The presence of these thick, refractory and anhydrous peridotite residues beneath Archaean crustal regions is widely held responsible for the inherent stability of Archaean cratons (Pollack 1986; Durrheim & Mooney 1994; Griffin et al. 2003; Kaban et al. 2003). The Archaean North China Craton (NCC), however, is an anomaly. Whereas it formed in the Archaean, complete with a thick, diamondbearing lithospheric keel, and behaved as a stable lithospheric block throughout the Proterozoic and Palaeozoic, this stability ended abruptly

during the Mesozoic, when the Eastern Block of the craton became reactivated, as witnessed by voluminous magmatism and widespread deformation (Liu 1992; Li & Yuan 2001; Xu et al. 2002; Meng 2003; Zhang, H. F. et al. 2003; Huang & Zhao 2004; Wu et al. 2005; Xu et al. 1991). Today, the Eastern Block of the craton has anomalously high heat flow and is seismically and volcanically active (six of the 10 deadliest earthquakes in recorded history have occurred in the eastern North China Craton; Kusky 2003; Zhang, Y. Q. et al. 2003). Estimates of lithospheric thickness are less than 100 km (Ma et al. 1984; Ma & Wu 1987; Yuan 1996; Yang 2003) and mantle xenoliths carried by Cenozoic basalts have relatively fertile bulk compositions (Menzies et al. 1993; Xu et al. 1998; Griffin et al. 1998; Zheng et al. 1998; 2001; Xu et al. 2002; Wu et al. 2003a, b; Lee & Walker 2006), contrasting markedly with the refractory peridotites characteristic of Archaean cratons (and found in Palaeozoic kimberlites from this very region) and the evidence for thick lithosphere preserved in the diamonds. Thus, the Eastern Block of the NCC is no longer a craton (see Kusky et al. 2007), and it appears that more than 150 km of mantle keel was removed from the base of the lithosphere sometime in the Mesozoic.

LITHOSPHERIC THINNING IN EASTERN ASIA

A variety of types of data indicate that the eastern half of the subcontinental lithospheric mantle of the North China Craton has been dramatically thinned or lost during the Mesozoic. These include geophysical surveys of gravity, magnetic and seismic data for deep crustal structure, geochemical and petrological data from magmas and entrained xenoliths, and other less direct indicators such as changes in topography, regional tectonics, and basin and thermal evolution. Many of these aspects of the root loss have been discussed elsewhere in this volume, and we summarize the more salient features here. The north –south gravity lineament is a major gradient in Bouguer gravity anomalies that corresponds roughly to the border between the Eastern and Western Blocks (or areas with and without

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root loss), but also extends north and south for thousands of kilometres past the borders of the NCC (Fig. 2). As the gravity lineament also corresponds to a major elevation change, areas of Tertiary basin formation along major faults, and a change in mantle seismic velocity at 150– 200 km depth (e.g. Niu 2005), it may represent a major crustal structure that parallels the Pacific subduction margin. The region to the east of the lineament has experienced Late Mesozoic– Tertiary extension, whereas the area to the west generally has not and has thicker crust. The north– south gravity lineament is interesting also because it bounds areas that to the west have thick crust and 150– 200 km thick lithosphere, large negative Bouguer anomalies, and low heat flow. Sub-Moho seismic Vp velocities west of the lineament are high, with values ranging from c. 8.1 to 8.3 km s21 (Fig. 3). However, to the east the crust and lithosphere are generally thinner, there is high heat flow, and the

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Fig. 3. Map showing depth to the low-velocity zone (modified after Griffin et al. 1998). NSGL, north– south gravity lineament.

regional Bouguer anomalies are zero to slightly positive. Sub-Moho seismic velocities are lower than to the west, ranging from 7.6 to 7.7 km s21, with some faster regions (implying partial root loss?). Tomographic profiles from the Eastern Block (Yuan 1996) show a ‘blobby’ velocity structure for the lower lithosphere, suggesting only partial root loss. The Eastern Block is seismically very active, experiencing many earthquakes of magnitude 8, which include six of the 10 most deadly events in recorded history (Kusky 2003) that left more than one million people dead in total. Whatever the process of root loss, it appears to have caused continuing lithospheric instability that is currently poorly understood. Understanding root loss may shed light on this active seismicity, potentially saving numerous lives. Although extending for thousands of kilometres along the Pacific rim, Mesozoic granitoids and gold deposits form a ring around the Eastern Block, perhaps reflecting some interaction between subduction-related processes and loss of the lithospheric root (Fig. 4). The granitoids, and associated

ore-bearing fluids, may contain one of the best and most detailed records of the history of root loss beneath the NCC, and detailed study of these plutons may yield insights into the chemical and physical environments associated with foundering of subcrustal lithosphere and the mechanisms of root loss. Loss of the lithospheric root is also shown by the compositional data for mantle xenoliths entrained in Early Palaeozoic and Mesozoic to Tertiary kimberlites and volcanic rocks. The oldest kimberlites (450 –490 Ma) are the Fuxian and Mengyin pipes in the west (Fig. 5), whereas the Teiling intrusions are of Cretaceous to Tertiary age. Xenoliths from Nushan are only 0.5–0.8 Ma old, and these, together with the older examples, provide a 500 Ma history of mantle samples from beneath the NCC. Geotherms based on mantle xenolith data (Ryan et al. 1996; Griffin et al. 1998; Gao et al. 2006) and garnet concentrates show that in the Ordovician, the Eastern Block had a low conductive cratonic geotherm, with many samples coming from beneath the diamond stability field. The Ordovician lithosphere–asthenosphere boundary is estimated to have been at about 180 km depth (Yuan 1996; Griffin et al. 1998, 2003). In contrast, compositional data for the younger mantle samples reveal a high geotherm and a lithosphere–asthenosphere transition that had risen to about 80 km depth. Compositional data for xenoliths thus clearly show the loss of the lithospheric root beneath the eastern NCC, but do not yield information on exactly when this loss may have occurred, why it occurred, or what the loss means for cratonic evolution. Basalts erupted through the crust of the Eastern Block also show a change in composition from Mesozoic to Tertiary, with high-Mg andesites or adakites showing evidence for lower crustal foundering in the Jurassic–Cretaceous.

Understanding root loss timing and mechanisms It appears that a major difference in crustal thickness between the eastern and western parts of the North China Craton is coincident with the prominent north–south-striking topographic change and gravity anomaly (e.g. Griffin et al. 1998; Niu 2005). The present-day distribution of tectospheric root provides constraints on the timing and causes of the loss of the root under the eastern half of the craton. A review of the papers in this volume (and references therein) reveals several themes that may be relevant for understanding the loss of the root, as follows.

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Granites

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128

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and

Gold

42

n sha

lt

be

fau

lt

n Ya

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Ta n-

Lu

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Central Orogenic belt

Jiaodong (Shandong) Peninsula

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Qingdao

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Yellow Sea Qin orog glong - Da en bi

N

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0

n

Mesozoic granitoid

Explanation

km

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Mesozoic gold deposit

Fig. 4. Mesozoic gold and granite provinces of the NCC. (Note how the gold deposits and granites outline a ring around the Eastern Block of the craton (EB), suggesting that they may delineate the limits of the area of root loss.) Modified after Goldfarb et al. (2001), Hart et al. (2002), Zhou et al. (2002) and Wu et al. (2005). COB, Central Orogenic Belt.

(1) The timing of loss seems to be c. 140 –120 Ma, based on the change in character of magmatism, structural evolution, basin formation, and types of xenoliths entrained in magmas and kimberlites. (2) The timing of the loss is roughly coincident with and slightly younger than continent– continent collisions in the Dabie Shan–Su-Lu orogen in the south and along the Solonker suture in the north. (3) Before root loss, the sub-continental lithospheric mantle experienced considerable hydration from dehydration reaction above subducting slabs.

This might have caused significant hydrationrelated weakening of the sub-continental lithospheric mantle. (4) The root loss seems to be associated with a change from contraction to extensional tectonics in the upper crust; thus, root loss may have triggered upper crustal extension, or upper crustal extensional collapse of thickened orogens may have been the final trigger for root loss. (5) The boundary between parts of the subcontinental lithospheric mantle that were lost and were not lost is near the Archaean suture in the

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Fig. 5. Map of the eastern NCC showing distribution of kimberlites of different ages that contain mantle xenoliths.

Central Orogenic Belt, suggesting that some inherent difference in the root between the Eastern and Western Blocks may have led to the eastern

side being lost whereas the western side was retained. Alternatively, the eastern side of the root may have been more greatly affected by

LITHOSPHERIC THINNING IN EASTERN ASIA

(a) pre-weakening

(b) triggers

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(c) mechanisms of root loss - continental collision

NCC

NCC

NCC

NCC

-extension

hydro-weakening changes rheology by adding water from dehydration reactions in subducting slabs

NCC

density foundering

delamination

NCC

NCC

thermo-chemical erosion

Fig. 6. Model showing mechanisms of root loss, and possible triggers for these mechanisms. Hydro-weakening may aid any of the mechanisms. Delamination and foundering may be triggered by collision, extension, or strike-slip faulting, and may in turn cause thermo-chemical erosion. Plume impingement may also independently cause thermo-chemical erosion.

Palaeopacific-related processes such as greater hydration (from the Pacific), leading to its preferential loss.

Hypotheses Various hypotheses have been advanced in attempts to explain the cause of the loss of the lithospheric root beneath the NCC; these can be broken down into four general physical categories (Fig. 6), as follows. (1) Mechanical disruption: this includes collision and rifting, both active and passive. (2) Thermo-chemical erosion: upwelling asthenosphere in the mantle wedge erodes lithospheric keel incrementally. (3) Density foundering or delamination: sinking of tectonically thickened lithosphere into convecting mantle. (4) Hydro-weakening of the subcontinental lithospheric mantle by multiple subduction events pre-weakens root for delamination during younger collisions; hydration also lowers the melting temperature of the root, allowing easier extraction of a silicic melt phase, and thereby increasing the density of the root and favouring delamination. Data presented in the papers in this volume provide some clues that help distinguish between these types of physical models. Understanding the timing of root loss and correlating this with regional tectonic and global (mantle) events helps differentiate between specific triggering mechanisms, and

provides tests for further work. For instance, mechanical disruption of the lithospheric root, by collisional, strike-slip or other mechanical processes, should lead to the development of steep physical gradients related to tectonic boundaries, which themselves should show a relationship to surface tectonic features. In contrast, thermo-chemical erosion, if related to upwelling asthenosphere in the mantle wedge, may erode the mantle root incrementally, forming smooth boundaries. If the thermochemical erosion is related to arc- or slab-derived fluids, associated magmas should possess some subduction-like geochemical signature, and show a geometric relationship to subduction-zone, shallow gradients that may cut across crustal structures. However, if the thermo-chemical erosion were related to plume impingement (Wilde et al. 2003; Deng et al. 2004), we would expect sharp lithospheric boundaries to be truncated and smoothed by longer-wavelength plume-related mantle. These trends should be resolvable with future seismic, petrological and structural studies in the NCC. Density foundering, or ‘delamination’ of tectonically thickened lithosphere might be sudden and simultaneous across a region the size of the Eastern Block of the NCC (e.g. Schott & Schmeling 1998). Such catastrophic loss of the mantle root should be detectable by the almost simultaneous responses across the eastern half of the craton, a pattern that has not yet been established or tested for. Mantle weakening by hydration is able to work with the other mechanisms described above and

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contribute to the ability of the sub-continental lithospheric mantle to delaminate or be lost by other mechanisms (Fig. 6). When greenschist-grade oceanic lithosphere subducts, it hydrates the upper mantle. In the Western Pacific, where 90% of the world’s small ocean basins are concentrated, subduction from the Pacific (westwards) combined with Indo-Australian (northwards) plates has doubled the amount of water transported to the mantle. The dual subduction of the Pacific and Indo-Australian plates has created a double-sided, triangular or Y-shaped zone, implying that double the amount of water has been transported into the mantle for a total at least 700 Ma, the combined length of time of subduction from the Pacific and Indo-Australian plates (Maruyama et al. 2004). Therefore double subduction allowed twice the amount of water to reach the transition zone under the marginal basins of the Western Pacific and under the North China Craton. Besides being subducted by the Pacific plate from the east and the Indo-Australian plate from the south, the Eastern Block of the North China Craton was affected by other subduction zones. The craton is surrounded by sutures that are the sites of former oceans that were subducted in all cases towards and under it; the Solonker suture on the immediate northern side, the Mongol–Okhotsk suture farther north in eastern Mongolia, the Qinling –Dabie Shan suture on the immediate southern side, and the Song Ma suture farther south in Indo-China. The water in oceanic basalts is transported by subduction into the mantle transition zone and stored in nominally anhydrous minerals in hydrogen-related point defects, if the geotherm is cold enough (Maruyama 2003). Wadsleyite (b-Mg2SiO4; stable below 350 km) and ringwoodite (g-dimorph of forsterite), the principal components of the transition zone, can contain up to 3.3 wt% and 2.2 wt% H2O respectively (Smyth et al. 2003). If saturated, a transition zone containing 70 modal % wadsleyite could contain four times the amount of water in the Earth’s hydrosphere, and would form the largest reservoir of hydrogen in the planet (Smyth & Frost 2002). Olivine, generally considered to be the most abundant mineral phase in the upper 400 km of the mantle, can store up to 2000 ppm by weight of H2O at 13 GPa and 1100 8C (Smyth et al. 2003). The major cause of the 410 km seismic discontinuity at the top of the transition zone is believed to be the reaction of olivine to modified spinel, and modified spinel can store up to 3.1 wt% H2O at 15.5 GPa (Inoue et al. 1995). Thus it is generally considered that water in a subducting slab would react with olivine to form hydrous modified spinel that would transport considerable amounts of water into the transition zone (Inoue et al. 1995).

This mantle hydration lowered the mantle viscosity, and lowered the melting temperature of the mantle to promote magmatic activity. Perhaps the hydro-weakening of the upper mantle has allowed fragmentation of the oceanic lithosphere and the formation of multiple micro-plates. The small ocean basin basalts are 0.2 wt% richer in water than normal mid-ocean ridge basalt (MORB); this lowers the solidus temperature by 150 8C. At 200 km depth addition of 0.2% water lowers the melting temperature (solidus) by 5008 (Maruyama et al. 2004). Only 100–1000 ppm addition of water decreases mantle viscosity by about two orders of magnitude (Maruyama et al. 2004; Niu 2005; Windley et al. 2005). Stored water in hydrous wadsleyite and ringwoodite at 410–660 km depth would become unstable with time as a result of conductive heating from underlying lower mantle (Maruyama et al. 2004; Niu 2005; Windley et al. 2005), and would eventually decompose into anhydrous phases to release free water, which would result in mantle melting and hydrous plumes. This was possibly the principal cause of the fragmentation of the oceanic lithosphere in the Western Pacific. Hydro-weakening by the cumulative effects of subduction has the potential to work with some of the other mechanisms described above to contribute to root loss beneath the NCC. The sub-continental lithosphere under Eastern China has been subjected to subduction of both the Pacific and the Indo-Australian plates for c. 150 Ma and c. 550 Ma, respectively, but also to northwards subduction on the site of the Dabie Shan suture for c. 500 Ma (Sengo¨r & Natal’in 1996; Ratschbacher et al. 2003; Oh & Kusky 2006) and to southwards subduction on the site of the Solonker suture (in Inner Mongolia) for c. 250 Ma (Xiao et al. 2003). The total length of slabs subducted under Eastern China in the last 150 Ma is c. 18 000 km, which is more than twice that subducted under any other area of the circum-Pacific. Clearly more oceanic lithosphere has been subducted under Eastern China than under any other part of the continents; this probably contributed to the ability of the NCC to lose its root. Accordingly, the multiple subduction must have excessively hydro-weakened the upper mantle under Eastern China. However, it did not cause the break-up of the continental lithosphere into several micro-plates. In contrast, the hydro-weakened mantle, being the keel of the eastern half of the Archaean North China Craton, apparently responded by delamination. A question remains: was there a tectonic mechanism that triggered the delamination? Considering that the most likely time of delamination was in the Cretaceous, tectonic forcing by orogenic collapse at the southern and northern ends of the delaminated zone provides a viable mechanism. Both the

LITHOSPHERIC THINNING IN EASTERN ASIA

Dabie Shan and Solonker orogens underwent massive Himalayan-style collisions, and intense post-collisional thrusting during the Jurassic (Hacker et al. 2000; Xiao et al. 2003; Li et al. 2006). Extensional collapse of the thrust-thickened lithosphere of these orogens could have triggered the delamination of the hydro-weakened lithosphere situated between them. In the case of the Solonker orogen the thrusting was accompanied by widespread granitic magmatism in the Jurassic, which was followed by the formation of many metamorphic core complexes in the Cretaceous, and crustal-scale extension in Eastern China gave rise to the many sedimentary basins, including the North China basins on top of the delaminated zone in the eastern North China Craton. If this were the mechanism responsible for the delamination in the Cretaceous, new fertile asthenosphere would have risen into the foundered zone of the upper mantle, and would have given rise to many alkali basalts and possibly some barren kimberlites in the Cenozoic.

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main characteristics that should produce features that are resolvable with future studies. Table 1 shows these hypotheses and their likely manifestation in terms of observables. The main observable features of the sub-continental lithospheric mantle will be seismic properties, mantle samples from xenoliths, melts and fluids from mantle-derived magmas, and surface structural and topographic responses to changing mantle thermal and density conditions. Elsewhere in this volume (Kusky et al. 2007) we outlined some of the main tectonic events in the crust that may relate to formation and loss of the root, and other papers have described detailed aspects of igneous, sedimentary, structural, hydrothermal and tectonic processes that may be related to root loss, and geophysical tests of the deep structure. Further work is needed to test these links, and it is clear that an integrated multidisciplinary programme is needed to further understand root loss beneath the NCC.

Conclusions Hypothesis testing Below, we propose tests of these and other related hypotheses, by dividing the models into their

The North China Craton is the best-known place in the world where a craton grew a sub-continental lithospheric root in the Archaean then lost part of

Table 1. Specific hypotheses that have been advanced to explain root loss, and their tests Model and proponents Regional extension and rifting caused mechanical dispersal of old, cold mantle and caused upwelling of new hot asthenosphere (Yuan 1995; O’Reilly et al. 2001) Convective overturn (upwelling?) accompanying Jurassic –Tertiary subduction of Pacific plate (Griffin et al. 1998); subduction of the Pacific plate caused plume-like upwelling (Deng & You 1985) Collision of NCC and Yangtze Cratons (Griffin et al. 1998; Gao et al. 2002), Solonker block and northern margin of NCC Subduction under Dabie Shan and Mongol– Okhotsk regions caused destabilization, thinning and replacement of lithospheric mantle in the Mesozoic followed by new underplating in the Cenozoic (Zhang, H. F. et al. 2003) Active rifting as a result of plume impingement Thinning during Gondwana fragmentation, subduction, mantle overturn, downwelling and mantle delamination (Wilde et al. 2003) Collision of India and Asia (Menzies et al. 1993)

Tests of model New lithosphere would be present only beneath rifts, and ancient root would still be present as intra-rift ‘islands’; these ancient relicts may be detected from both xenolith and seismic tomographic investigations Root loss should occur in zones related to subduction geometry

Timing should be shortly after collision This mechanism may work independently, or may ‘pre-weaken’ mantle root preparing it for loss in later events such as collisions

If plume related, gradients will be smooth and may cut across all tectonic boundaries Boundaries should be diffuse between lost and retained keel; timing should correlate with break-up; chemistry of basalts should be plume-like Timing should be post-collision

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this root in Phanerozoic tectonism (Marotta et al. 1998; Meisner & Mooney 1998). Understanding of the processes of root formation and loss is in its infancy, yet some salient points appear relevant. First, the root of the North China Craton was apparently formed during initial formation of the craton between 3.8 and 2.5 Ga, with roots of different character forming beneath the Eastern and Western Blocks, which collided at 2.5 Ga (Kusky 2001, 2004; Li & Kusky 2006). A major continent–continent collision at 1.85 Ga on the north margin of the craton significantly modified the root, and locally replaced part with Paleoproterozoic mantle (Kusky & Li 2003). The North China Craton was surrounded by subduction zones for much of the Palaeozoic and Mesozoic. Dehydration reactions in these slabs released fluids that significantly hydrated the mantle above the subduction zones (the root of the NCC), changing the rheology of the mantle root and causing significant hydration-related weakening. The idea that subduction of water into the mantle caused hydro-weakening of the subcontinental lithosphere and was responsible for the thinning–delamination under the Eastern block of the North China Craton came independently from Niu (2005) and Windley et al. (2005). However, whereas Niu (2005) considered that subduction by the Pacific plate was sufficient to carry water to the upper mantle, Windley et al. (2005), building on the ideas of Maruyama et al. (2004) and Komiya & Maruyama (2006) of double subduction, as summarized above, extended the process to include the subduction zones sited on the Solonker, Dabie Shan and Mongol –Okhotsk sutures. Some Western Pacific marginal basins physically link with onshore areas in Eastern China underlain by thinned lithosphere, demonstrating a probable genetic relationship between the processes of hydration of the mantle transition zone responsible for formation of the marginal basins, as outlined above, and the processes that gave rise to the thinning of sub-continental lithosphere. The most prominent example is the Bohai Basin, the initial extension of which began in the Palaeocene, probably triggered by subduction roll-back of the oceanic Pacific plate (Allen et al. 1997). Extension continued until the end of the Oligocene (Zhao & Zheng 2005). Within Eastern China the Bohai Basin is an area with thickest Cenozoic sediments (4000–7000 m; Allen et al. 1997), the highest surface heat flow 1.8 –2.0 heat-flow units (hfu), Ma 1987), and thinnest lithosphere (50–60 km; Ma 1987). Major continent–continent collisions on the southern and northern margins of the craton in the Mesozoic may have mechanically triggered the loss of the hydration-weakened lithospheric

root, triggering massive replacement of the sub-continental lithospheric mantle with new fertile mantle in the late Mesozoic and Cenozoic. The tectonic style in the crust changed at this time, from contraction to extension, and the processes of root loss and tectonic regime seem intimately related, although cause and effect have yet to be separated. We propose that post-collision thrusting in the Jurassic in the Dabie, Solonker and Mongol– Okhotsk orogens triggered collapse by the early Cretaceous of the hydro-weakened and eclogitized crustal roots of these orogens, which in turn triggered delamination of the intervening subcontinental lithosphere that itself had been hydroweakened. Replacement, either physically or chemically, by fertile asthenosphere in the Cenozoic led to extrusion of extensive alkali flood basalts throughout the Eastern Block of the North China Craton and in the Songliao Basin to the north. We thank our many colleagues who have worked with us in the North China Craton, and provided stimulating discussions about the interpretation of regional tectonics. We especially acknowledge the contributions of J. H. Li, R. Rudnick, B. Mitchell, S. Z. Li, A. Polat, L. Wang, A. Kroner, G. Davis, F. Y. Wu, S. Wilde, W. J. Xiao and J. H. Guo. Reviews by R. Goldfarb and A. Polat greatly improved the manuscript. This work was supported by Chinese Academy of Sciences grant KZCX1-07 awarded to M.Z. and Rixiang Zhu, the US NSF grants 01-25925 and 02-07886 awarded to T.M.K, by St. Louis University, the Chinese Academy of Sciences, Peking University, University of Leicester and Ocean University of China.

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Index Page numbers in italics refer to Figures; page numbers in bold refer to Tables Aishan granitoid, 303 Alashan Block, 3 Alpine-Himalaya orogen, 2 Altaides, 224, 225 Altyn Tagh, 9 Altyn–Tagh fault, 241 amphibolite facies, Central Orogenic Belt, 4 40 Ar/39Ar dating Dandong granite, 176, 180 extension structures, 157, 162 Inner Mongolian orogen, 227 Jiaodong Peninsula, 312 Mesozoic igneous rocks, 81, 104, 195, 202, 203, 205, 206, 217 Archaean cratons characteristics, 331–332 map, 332 Bainaimiao island arc, 10, 225, 227, 229– 230 Baiqi Formation, 194 Banded Ironstone Formation, 5 Baoerhantu arc, 9 Baolidao complex, 10 Beihuaiyang fold belt, 241 Beijing graben, 192 Beijing– Chengde basins, 243 Beipiao Formation, 243 Bohai Sea basins, 16– 19 Bouguer gravity, 20 map of China, 19, 333 North China Craton, 132 Yellow Sea, 283–284, 286 Cambrian events on North China Craton, 9, 225, 234 carbonatites, 57, 61 Carboniferous events on North China Craton, 9, 234 Cathaysia Craton, 2, 289 Cenozoic events on North China Craton active structures, 18 crustal extension, 16– 19, 48 kimberlites and xenoliths, 12, 21, 46, 334, 336 lithosphere evolution, 70, 77–78 Central Asia Fold Belt, 241 Central Asia orogen, 2, 36, 56, 78 Central Asian Orogenic Belt (CAOB), 95, 132 Central China orogen, 1, 2 Central Orogenic Belt, 3, 4, 7 Changping fault, 192 Changzhougou Formation, 190 Chaoyang Basin, 203 Chengde Basin, 203 Chengde metamorphic complex, 6 Chengde thrust fault, 192 Chengde– Xinlong thrust, 176 Chifeng basalt, 37 Chifeng Basin, 203 Chifeng Bayan Obo fault, 10 coesite in eclogite, 36

collision events on North China Craton Precambrian, 7 –8 Triassic, 180, 285– 288 compression events East Laioning Peninsula, 163– 164, 166 Yanshanian, 322 Cretaceous events on North China Craton, 12, 37, 38, 48, 70, 124 extension, 181– 183, 205– 207 basins, 202 fill and facies, 209–214, 246, 248– 252 migration, 255– 256, 259 rift-type, 193, 196 drivers, 217 –218 East Laioning Peninsula case study, 166–167 central area, 155 –157 north area, 157–158 south area, 158–163 structures high strain, 214– 216 low strain, 207–209 timing of, 216– 217 inversion, 216 melting, modelling of, 182 metamorphic core complexes, 202 stratigraphy, 211, 213 tectonism and sedimentation, 204– 205 crust contribution to SCLM geochemical evidence, 63–67 mechanisms, 67– 68 modelling behaviour, 68– 71 imaging. See seismic tomography structure for North China Craton, Eastern Block, 277 thickening, Yanshanian, 318, 322– 325 Dabeigou Formation, 194, 196, 243, 246, 247, 253, 255 Dabie Mountains Orogenic Belt, 59, 240, 241 Dabie UHP belt, 132 Mesozoic magmatism chemical data, 139–140 isotope data, 142–143 rare earth element data, 144 Dabie –Sulu suture zone, 171 Dahongyu Formation, 190 Dahuanggou pluton, 173 Dandong granite deformation, 175–176 timing of, 176– 180 Dashi Formation, 243, 253 Daye Formation, 253 deformation studies. See also compression extension delamination North China Craton lithosphere, 218–219, 317, 326 –328, 337 –339 Devonian events on North China Craton, 9, 225, 234 diamond in eclogite, 36 See also kimberlite Diaojishan Formation, 37

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Don Ujimqin arc, 9 Dong Ujumqi magmatic arc, 225 Donggan gabbro, 37 Dongjiagou shear zone, 179 Donglanggou Formation, 195, 196 Donglingtai Formation, 37 Dongqin pluton, 173 Dongqiyishan arc, 9 Dongye Group, 5 Doucun Group, 5 earthquake epicentres, North China Craton, 295 East Liaoning Peninsula compression, 163–164, 165 summary of evidence, 166 extension Buyunshan area, 157 Liaoyang– Benxi area, 157–158 southern peninsula area, 158–163 summary of evidence, 166 –167 Xiuyan area, 155– 157 geological setting, 154–155 tectonic studies East Marginal Fault of the Yellow Sea (EMFYS) 282, 283, 285, 289 East Tianshan Block, 241 Eastern Block of North China Craton. See North China Craton, Eastern Block eclogite, 36 Enshoo Terrane, 10 Eren Hot–Hegenshan ophiolite zone, 225, 226, 227, 232 Erlian Basin, 202, 216 extension, 181–183, 205–207 basins, 202 fill and facies, 209– 214, 246, 248–252 migration, 255–256, 259 rift-type, 193, 196 drivers, 217– 218 East Laioning Peninsula case study, 166– 167 central area, 155– 157 north area, 157– 158 south area, 158 –163 structures high strain, 214–216 low strain, 207– 209 timing of, 216–217 Fangcheng basalts, 37, 38–39 Fanghushan Formation, 243, 248, 253 Fangshan metamorphic core complex, 172 Farallon Palte, 217 Fenghuangtai Formation, 243, 248, 249, 250, 253, 256 Fengning –Weichang basins, 243 Fengzoshan Group, 303 fluid inclusion analysis gold from Jiaodong Peninsula, 306 microthermometry, 308–309 petrography, 306–308 Foziling Group, 241, 248 Fuping Complex, 5 Fushan complex, 37, 38 Fuxin basalt, 37 Fuxin Basin, 203, 215 Fuxin Formation, 38, 204, 243, 246, 253

Gaofan Group, 5 geochemistry analytical methods, 105, 112 Mesozoic magmatism, 80, 83, 84, 85 enrichment processes, 94– 95 evidence of SCLM, 63–68, 86–88 major elements, 88– 90, 106–111, 112, 134–142 Sr-Nd isotopes, 92–94, 112, 114, 115, 116, 142–143 trace elements, 61– 63, 80, 90–91, 101, 106– 111, 112, 134– 142, 324, 325 mantle-derived igneous rocks of North China Craton evidence for crustal contribution, 63– 67 mechanism for crustal contribution, 67–68 modelling behaviour, 68– 71 Mesozoic volcanics, 260 geochronology Dandong granite, 175– 176, 177– 178, 179 mafic magmatism, 79– 81 gneiss, 36 Gobi Basin, 202, 216 gold, 22, 23, 294, 300 Jiaodong Peninsula deposit classification, 304–305 fluid inclusions, 306 microthermometry, 308– 309 petrography, 306– 308 geological setting, 303–304 hydrothermal history, 310 sources, 311–312 tectonic setting, 313–314 timing, 310–311 paragenesis, 305–306 Gonganzhai Formation, 243, 253 granulites Central Orogenic Belt, 4 Henshan high-pressure granulite, 6 North China Block, 78 greenschist facies, Cental Orogenic Belt, 4 Guangminshan pluton, 173 Gudaoling granite, 173 Guojiadian pluton, 174, 180 Guojialing granitoid, 176, 303 Guojialing pluton, 180 Guojiazhai Group, 5 Guyi gabbro, 37 Guzheng complex, 38 Gyeonggi massif, 281 Haifanggou Formation, 243 Hailar Basin, 202 Hanshan microcontinent, 9 Hanxing complex, 39 heat flow history. See vitrinite reflectance Hefei Basin, 240 –241, 243, 246– 252, 256– 259 Hegenshan ocean, 225 Hegenshan ophiolite, 225 Hegenshan suture, 9 Heishidu Formation, 243, 248, 251, 253 Hengshan complex, 10 Hengshan high pressure granulite, 6 Hengshan metamorphic complex, 6 Hengshan Plateau, 3

INDEX high field strength elements (HFSE), 56, 61– 63, 78, 79, 101, 324, 325 Hohhot metamorphic core complex, 202 Honam shear zone, 281 Hongshan complex, 38 Hongshishan suture, 9 Houcheng Formation, 37, 192, 193, 195, 196 HREE. See rare earth elements (REE) Huaian metamorphic complex, 6 Huailai Basin, 240 Huajiahu Formation, 243, 253 Huajiying Formation, 194 Huaqiying Formation, 37 Huaxia Block, 241 Hungshandian fault, 192 Hutag Uul Terrane, 10 igneous processes. See magmatism Imjingang fault belt (IFB), 282 Imjingang Group, 282 Imjingang orogen, 281 Indo-Sinian orogeny, 12 Inner Mongolia–Daxinganling Orogenic Belt (IMDOB) geological setting, 225– 226 history of research, 223–225 lithospheric thinning, 234 –235 ophiolite zone geochronology Bainaimiao arc, 227, 229 –230 Eren Hot– Hegenshan, 231 Jiaoqier– Xilinhot, 230–231 Ondor Sum–Xar Moron, 226– 229 Solon Obo–Linxi, 230 Sunid–Baolidao arc, 231–232 ophiolite zones, 225 Eren Hot– Hegenshan, 226 Jiaoqier– Xilinhot, 226 Ondor Sum–Xar Moron, 225 Solon Obo–Linxi, 225 tectonic evolution, 233– 234 timing of suturing, 232–233 Inner Mongolian Suture Zone. See Inner Mongolia– Daxinganling Orogenic Belt isotope signatures and dating. See 40Ar/39Ar; K-Ar; Pb/Pb; Rb/Sr; Re-Os; Sm-Nd; 87Sr/86Sr; Sr-Nd; U-Pb Izanagi Plate, subduction, 183 Jaio-Liao massif geological setting, 173– 175 map, 172 relation to Sulu orogen, 180– 181 remelting evidence, 182 Jialingjiang Formation, 253 Jianchang basalt, 37 Jianchang Basin, 203, 209– 212, 240 Jianghan Basin, 243 Jiaodong Group, 303 Jiaodong Peninsula, 132, 174 crustal age, 180 geological setting, 173 Mesozoic magmatism chemical data, 137– 138 isotope data, 142– 143 rare earth element data, 144

347

metallogenesis in the gold province, 303–304 deposit classification, 304–305 fluid inclusions, 306 microthermometry, 308– 309 petrography, 306– 308 hydrothermal history, 310 sources, 311–312 tectonic setting, 313– 314 timing, 310– 311 map, 304 paragenesis, 305–306 Jiaoliao Block, 3 Jiaoqier– Xilinhot ophiolite zone, 225, 226, 227, 231–232 Jimo volcanics, 37, 38–39 Jinan gabbro complex, 37, 38, 39 Jinan norite, 80 Jining Longhua fault, 10 Jinshan Group, 303 Jinshan suture, 241 Jinzhai–Huoshan–Shucheng fault, 242 Jiufotang Basin, 203, 212– 214 Jiufotang Formation, 204, 211, 213, 243, 246, 253, 255 Jiulongshan Formation, 195, 243, 244, 245, 253, 255 Junggar-Hegen suture, 204 Jurassic events on North China Craton, 12, 37– 38, 48, 70, 79, 124 tectonics and sediments, 202–204, 204–205 K-Ar dating, 173 Inner Mongolian orogen, 226, 227 mantle and lower crustal xenoliths, 58, 59 Mesozoic igneous rocks, 81, 104 kimberlites, 331 Cenozoic, 12, 21, 334, 336 Mesozoic, 12, 46, 21, 334, 336 Palaeozoic, 21, 35, 55, 62, 77, 78, 131, 201, 267, 334, 336 Korean Peninsula, 283, 288 Kuancheng Basin, 240 Kunlun-Qilian-Qinling fault-fold system, 241, 283 Kunyushan granitoid, 303 Laioxi basins, 243 Laiwu gabbro-diorite complex, 37, 38, 39 Laiyuan complex, 37, 38, 39 Lanqi Formation, 38, 243 Laohugou Formation, 243 large ion lithophile elements (LILE), 56, 61–63, 78, 79, 101, 324, 325 Lhasa Block, 241 Liangbin pluton, 173 Liaodong Peninsula Dandong granite deformation, 175–176 timing of, 176– 180 geological setting, 173, 174– 175, 181 Triassic compression-Cretaceous extension evidence, 181 –183 Liaohe Group, 173, 174 Liaonan metamorphic core complex, 172, 180, 202, 203 Liaoning Province. See Eastern Liaoning Peninsula also under Yanshan orogenic belt

348

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Liaoxi basalt, 39 Linglong granite, 174, 176, 180, 303 Lingxiang Formation, 243, 253 lithosphere delamination, 218– 219, 317, 326–328, 337–339 evolution evidence in Late Mesozoic, 86–88 enrichment processes, 94–95 major elements, 88–90 Sr-Nd, 92– 94 thinning, 95–96, 234– 235, 333– 334 mechanisms, 45– 47, 317, 337 –339 model, 47–49, 68– 71 timing, 334– 337 trace elements, 90–91 Liuan– Qieshan fault, 241 Louzidian normal fault, 203 low-velocity layer (LVL), 20, 333 map of China, 334 LREE. See rare earth elements (REE) Luanjiahe granitoid, 303 Luanping Basin, 192, 203 Luanping– Chengde Basin, 240 Luxi basalt, 39 magmatism on North China Craton, 58– 59, 132 distribution, 37– 38 geochemistry, 80, 83, 84, 85 enrichment processes, 94–95 evidence of SCLM, 63– 68, 86–88 major elements, 88–90, 106 –111, 112, 134– 142 Sr-Nd isotopes, 92–94, 112, 114, 115, 116, 142– 143 trace elements, 61–63, 80, 90– 91, 101, 106–111, 112, 134–142, 324, 325 geochronology, 79–80, 81 isotope composition, 39 map, 79 mafic enclaves, 132– 134 mantle sources, 147 –148 origins, 143– 147 petrogenetic model, 148 petrography, 82 petrology, 38– 39, 80 Majiashan Formation, 243, 253 Manchurides, 224, 225 mantle, lithospheric, 55– 56 IOB, 55, 77 MORB, 55, 77 subcontinental, 77 chemical evidence for, 86–88 enrichment processes, 95– 96 major elements, 88– 90 Sr-Nd, 92–94 trace elements, 90– 91 subduction related, 77 mantle xenoliths, 21, 43, 44, 46, 58, 59, 131, 201, 267, 331, 334, 336 Maotanchang Formation, 243, 248, 249, 250, 253 Meichan Group, 241 Mengyin porphyries, 38–39 Mengyin shoshonite, 37 Mesozoic events on North China Craton magmatism, 58–59, 132 distribution, 37–38

geochemistry, 80, 83, 84, 85 enrichment processes, 94–95 evidence of SCLM, 63– 68, 86–88 major elements, 88–90, 106–111, 112, 134 –142 Sr-Nd isotopes, 92– 94, 112, 114, 115, 116, 142– 143 trace elements, 61–63, 80, 90– 91, 101, 106 –111, 112, 134– 142, 324, 325 geochronology, 79– 80, 81 isotope composition, 39 map, 79 mafic enclaves, 132– 134 mantle sources, 147–148 origins, 143– 147 petrogenetic model, 148 petrography, 82 petrology, 38– 39, 80 sedimentary basins, 240 –241 basin geodynamics, 261– 262 northern basins fill Cretaceous, 246 Jurassic, 242– 246 northern basins migration Cretaceous, 255–256 Jurassic, 252– 255 southern basins fill Cretaceous, 248–252 Jurassic, 248 southern basins migration Cretaceous, 256–259 Jurassic, 256 structural evolution, 259– 261 See also Cretaceous; Jurassic; Triassic metamorphic core complexes, 6, 172, 180, 202, 203 metamorphic facies, 4, 6, 78 metamorphism, UHP, 56, 303, 304 Mianlue suture, 11 microthermometry, gold from Jiaodong Peninsula, 308 –309 migmatization, Central Orogenic Belt, 4 Mishan fault, 304 Moho, depth at Yellow Sea, 285 Mongol–Okhotsk Ocean, closure, 70, 124, 181, 182 Mongolia Fold Belt, 241 Mongolian Block (microcontinent), 70, 124 Mongolian Orogen, 240 Nandaling Formation, 37, 45, 193, 195, 196, 243, 245, 253, 260 Nandazhai fault, 192 Nandian Formation, 194 Nantianmen Formation, 243, 246, 255 eNd Jiaodong Peninsula gold, 312 mantle derived rocks, 55, 56, 57, 60, 61–63, 65, 66, 67, 69 Mesozoic igneous rocks, 39, 80, 85, 86, 87, 88, 93, 105, 114, 115, 116, 145 Neimonides, 223 North Altyn Tagh suture, 9 North Boundary Fault, 293, 294 North China Basin, 202

INDEX North China Craton (Sino-Korean Block) Bouguer gravity, 19, 20 evolution summary, 25– 26 geological setting, 36 low-velocity zone depth, 20 tectonic divisions, 283 North China Craton, Eastern Block anomalous history, 332 crustal characteristics, 77, 277 geological setting, 36–37, 60– 61, 78–79, 201 Precambrian, 3 –5, 6, 7– 8 magmatism, 58– 59, 132 distribution, 37– 38 geochemistry, 80, 83, 84, 85 enrichment processes, 94–95 evidence of SCLM, 63– 68, 86– 88 major elements, 88–90, 106– 111, 112, 134– 142 Sr-Nd isotopes, 92–94, 112, 114, 115, 116, 142–143 trace elements, 61–63, 80, 90– 91, 101, 106– 111, 112, 134–142, 324, 325 geochronology, 79–80, 81 isotope composition, 39 map, 79 mafic enclaves, 132– 134 mantle sources, 147 –148 origins, 143– 147 petrogenetic model, 148 petrography, 82 petrology, 38– 39, 80 mantle state Cenozoic, 35–36 Mesozoic enrichment, 42 evolution, 40 heterogeneity, 39– 40 lithosphere thermal history reconstruction method, 270–274 results, 274–276 theory, 269–270 lithosphere thermal state, 267–269 lithosphere thinning mechanism, 45–47, 95–96 lithosphere thinning model, 47– 49, 68–70 lithosphere thinning timing, 42–45 olivine composition mapping, 41–42 Palaeozoic, 35 root loss evidence, 333 –334 mechanisms, 337–339 timing, 334–337 orogen–craton cyclicity, 20–24 sedimentary basins, 240– 241 basin geodynamics, 261–262 northern basins fill Cretaceous, 246 Jurassic, 242–246 northern basins migration Cretaceous, 255–256 Jurassic, 252–255 southern basins fill Cretaceous, 248–252 Jurassic, 248

349

southern basins migration Cretaceous, 256– 259 Jurassic, 256 structural evolution, 259– 261 seismic tomography, 20, 294, 295, 296 methods, 297 results, 297– 300 results discussed, 300 structural overview, 1– 3 tectonic evolution, 172, 239– 240 Cenozoic, 16– 20 Mesozoic, 16, 180 –183 Palaeozoic, 9, 12– 16 See also East Liaoning Peninsula; Yanshan orogenic belt North China intracontinental orogen. See Yanshan orogenic belt North Huaiyang basalt, 37 North Orthogneiss Unit, 172 North Qilian suture, 9 North–South gravity lineament (NSGL), 132 Northeast China fold belt, 102 Northern Dabie complex, 37 Nuhetdavaa Terrane, 10 olivine, composition mapping in mantle modelling, 41–42, 43–44, 45, 46 Ondor Sum complex, 10 Ondor Sum Ocean, 225, 235 Ondor Sum–Xar Moron ophiolite zone, 225, 226–229 ophiolites of Inner Mongolia –Daxinganling Orogenic Belt geochronology of zones Eren Hot–Hegenshan, 227, 232 Jiaoqier–Xilinhot, 227, 231–232 Ondor Sum– Xar Moron, 226–229 Solon Obo– Linxi, 227, 230 geological setting, 225 –226 Ordos Block, 201, 202 See also Western Block of North China Craton Ordovician events, 9, 225, 234 kimberlites, 21, 35, 55, 77, 78, 131, 201, 267, 334, 336 ore-bearing fluids, 334 See also gold osmium isotopes, 332 P waves. See seismic waves Pacific Plate (palaeo; Izanagi), 217 subduction, 47–49, 183– 184, 201 Palaeoasian (Palaeo-Asian or Turkestan) Ocean, 9, 12, 47–49 palaeogeography, Hefei Basin, 258 palaeotemperature analysis for North China Craton, Eastern Block. See vitrinite reflectance Pb/Pb isotope ratios mantle derived rocks, 55, 56, 57, 60, 61– 62, 64, 67, 71 Mesozoic igneous rocks, 39, 40 peridotites, 21, 43– 44, 62, 332, 336 See also olivine Permian events on North China Craton, 9, 225, 234 plagioclase, in mafic enclaves of Mesozoic igneous rocks, 133, 134 Puqi Formation, 243, 253

350

INDEX

Qaidam Block, 241 Qian Tang Block, 241 Qianshan granite, 173 Qianxi Basin, 240 Qiliashan fold belt, 241 Qingling –Dabie orogen, 11, 13, 56, 78, 79, 102 Qinglong foreland basin, 3, 6 Qingshan Formation, 38, 39, 45 Qingshan pluton, 180 Qingshilazi Formation, 194 Qinling orogen, 181 Qinling–Dabie collision zone, 281 Qinling–Dabie orogenic belt, 56 Qinling–Dabie– Sulu orogen, 172 Qinshila Formation, 243, 246, 253, 255 Queshan metamorphic core complex, 172 rare earth elements (REE), 61, 78, 80, 90, 94, 96, 101, 141– 142, 144 Rb/Sr dating, 94 Inner Mongolian–Daxinganling Orogenic Belt, 226, 227 Jiaodong Peninsula, 312 Liaodong Peninsula plutons, 173 mantle-lower crustal xenoliths, 58, 77 Mesozoic mafic igenous rocks, 81, 105, 115 Yanshan volcanics, 194–195 Re/Os ratio, 35 Red River fault, 241 Renao pluton, 173 rift basins, Cretaceous Dabeigou volcano-sedimentary cycle, 196 Nandaling volcano-sedimentary cycle, 193 Tiaojishan volcano-sedimentary cycle, 193, 196 Zhangjiakou volcano-sedimentary cycle, 196 Sanguliu pluton, 176, 179 Sangyi-Pinquan falt, 242 Sanjianpu Formation, 243, 248, 249, 250, 253, 256 sedimentary history relation to tectonism, 202– 205 sedimentary basins, 240–241 basin geodynamics, 261–262 northern basins fill Cretaceous, 246 Jurassic, 242–246 northern basins migration Cretaceous, 255– 256 Jurassic, 252–255 southern basins fill Cretaceous, 248– 252 Jurassic, 248 southern basins migration Cretaceous, 256– 259 Jurassic, 256 volcano-sedimentary cycles, 193, 196 seismic activity, 334 seismic tomography (high resolution), 20 data set, 294, 295, 296 results, 297 mid-lower crust, 298– 300 upper crust, 297–298 results discussed, 300 seismic wave velocities (vp), 20, 284, 287, 288, 333–334

Shahai Formation, 243, 246, 253 Shangdan suture, 11 Shanxi graben, 16–19, 19 Shanyi– Pingquan fault, 240 SHRIMP (sensitive high resolution ion microprobe), 6, 8 Dandong granite, 175– 180 Inner Mongolian–Daxinganling Orogenic Belt, 227, 228, 229 Jiaodong Peninsula, 132, 312 mantle-lower crustal xenoliths, 58–59, 324– 325 Shuangta pluton, 173 Siberian Craton (Plate), 70, 124, 181, 225 Silamulun suture. See Solonker Silurian events, 9, 225, 234 Sino-Korea-Yangtze collision zone, 282–283, 285–288 Sino-Korean Block (Craton). See North China Craton Sm/Nd isochron, 59, 77, 94, 227 Solon Obo–Linxi ophiolite zone, 225–226, 227, 230 Solon– Linxi suture, 9 Solonker (Suolun or Silamulun) suture, 8, 10, 12– 13, 204, 225 Song-Pan Block, 241 Songliao Basin, 202, 216 Songpan Ganzi orogen, 2 Sononker Suture, 181, 182 South Boundary Fault, 293, 294 South China Craton (Block), 70, 95 Mesozoic mantle-derived rocks, 59 tectonic divisions, 172 South China orogen, 13, 14 South Marginal Fault of Jeju Island (SMFJI), 282, 284, 285, 289 Southern Taihangshan metamorphic complex, 6 87 86 Sr/ Sr isotope ratios mantle derived rocks, 55, 56, 57, 60, 61– 62, 66, 67, 71 Mesozoic igneous rocks, 39, 40, 41, 80, 85, 86, 87, 88, 93, 94, 105, 114, 115, 117, 118, 145 subcontinental lithospheric mantle (SCLM), 131 Sulinheer Terrane, 10 Sulu orogen, 172 relation to Jiao-Liao massif, 180–181 Sulu orogenic belt, 56 Sulu suture belt, 282 Sulu UHPM, 303 Sunid Baolidao arc, 231– 232 Sunjiawan Formation, 243, 246, 253 Suolon suture. See Solonker Taihang Mountains, 3, 6 Taihang– Yanshan orogen, 132, 147 Mesozoic magmatism chemical data, 135–136 isotope data, 142–143, 145, 146 rare earth element data, 144 Taihanshan metamorphic core complex, 172 Tan-Lu (Tanlu) fault, 12, 56, 95, 132, 240, 241, 281, 293, 294, 302, 303, 304 Tanchen–Luzhou fault, 283 Tarim Block, 2, 13, 172, 241 tectonic divisions, North China Craton, 172 tectonic events, Mesozoic. See Jiao-Liao massif data thermal history reconstruction. See vitrinite reflectance thermometry, gold from Jiaodong Peninsula, 308–309 Tiajishan box fault, 191

INDEX Tian Shan–Yin Shan suture, 204 Tiaojishan Formation, 191, 192, 193, 195, 196, 243, 246, 247, 253, 255, 259, 260 Tien Shan-Xinganling fault-fold system, 283 tonalite–trondhjemite–granodiorite (TTG), 4, 36, 78 Tongjiapu tonalite, 173 Tongshi syenite, 37, 38, 42 trace elements, North China Craton Mesozoic mafics, 61–63, 80, 90– 91, 101, 106–111, 112, 134– 142, 324, 325 Trans-North China orogen, 4, 56 Triassic events in North China Craton, 12, 48, 164, 166, 234 collision, 70, 124, 180 compression, 181– 183 tectonics and sediments, 202– 204 tectonism/magmatism, 201 Tuchengzi Formation, 204, 205, 206, 243, 244, 246, 247, 253, 255, 259 Turkestan (Palaeoasian) Ocean, 9, 12, 47–49 U-Pb dating, 7, 8 Dandong granite, 175, 176, 177–178, 179 Inner Mongolian–Daxinganling Orogenic Belt, 228, 229, 230, 231, 232 Jiaodong Peninsula, 132, 312 mantle-lower crustal xenoliths, 58– 59, 68, 324– 325 Mesozoic igneous rocks, 81, 164 Yanshan volcanics, 194, 195 Uliastai margin, 10 ultra-high pressure metamorphic zone, 56 underplating, 49, 325– 326 vitrinite reflectance and palaeotemperature method of analysis, 270 results, 270–271, 272 results discussed erosion effects, 274 lithsphere boundary, 274–276 palaeo heat flow, 274 summary, 276– 278 thermal interpretation, 273– 274 theory, 269 volcanic events, Mesozoic, 319, 321 geochemistry analytical methods, 105, 112 results, 106–111 results discussed, 112– 116 North China Craton rocks compared, 120– 123 petrography, 103 –105 petrogenesis, 116 crustal contamination, 116 –117 fractional crystallization, 117–118 source involvement, 118– 120 tectonomagmatic significance, 123–125 volcano-sedimentary cycles Dabeigou, 196 Nandaling, 193 Tiaojishan, 193, 196 Zhangjiakou, 196 volcanogenic massive sulfide, 5 Wangshi Group, 38, 45 Waziyu metamorphic core complex, 202

351

Weideshan granitoid, 303 West Beijing Basin, 240 West Liaoning metamorphic complex, 6 West Tianshan Block, 241 Western Beijing basalt, 37 Western Block of North China Craton (Ordos Block), 3, 4, 36, 60 tectonic divisions, 172 Western Ordos Border Fault, 3 Western Shandong extension, 172 Wugingshan metamorphic core complex, 172 Wulian–Qingdao fault, 282 Wulian–Yantai fault, 172 Wulong pluton, 176, 179 Wutai Group, 5 Wutai Mountains, 3 Xar Moron fault, 10 xenoliths, 21, 43, 44, 46, 58, 59, 131, 201, 267, 331, 334, 336 See also kimberlite Xiadian Formation, 243, 246, 253, 255 Xiahuayuan Formation, 191, 193, 195, 243, 245, 253 Xiangdaopu Formation, 243, 252, 253 Xiangmo pluton, 173 Xiangxi Formation, 243, 253 Xi’anli complex, 38 Xiaoheishan pluton, 173 Xiaohungshan suture, 9 Xiaotian Basin, 241 Xiaotian Formation, 38, 243, 251, 252, 253 Xiaotian–Mozitan fault, 240, 241, 242 Xiguayuan Formation, 194 Xilamulun suture, 241 Xilin Hot microcontinent, 9 Xinglong fault, 192 Xinglonggou Formation, 37, 38, 260 Xingshikou Formation, 195 Xingzi metamorphic core complex, 172 Xinlonggou Formation, 243 Xinshikou Formation, 242, 243, 244, 245, 253 Xiong’er Group, 7, 8 Xishan basalt, 39 Xuanhua Basin, 240 Yagan– Onch Hayrhan metamorphic core complex, 202 Yangmulin pluton, 173 Yangtze Block. See Yangtze Craton Yangtze–Cathaysia Block collision zone, 283, 288– 289 Yangtze Craton (Block), 2, 36, 78, 132, 180, 181, 241, 281, 283, 289 Yangtze-Sino-Korea Block collision zone, 282 –283, 285–288 Yanlaio subsidence zone, 190 Yanshan fold-thrust belt, 202, 203, 204 Yanshan Movement (Mobilization), 239, 317 Yanshan orogenic belt, 103 block interactions, 191–193 compression evidence geochemistry, 322–324 xenoliths, 324 –325 delamination model, 326–328 fault distribution, 190

352

INDEX

Yanshan orogenic belt (Continued) geological setting, 189–190 lithosphere disturbance, 196–197 map, 102 Mesozoic intrusives, 320, 321 Mesozoic volcanics, 319, 321 geochemistry analytical methods, 105, 112 results, 106–111 results discussed, 112 –116 petrogenesis, 116 crustal contamination, 116–117 fractional crystallization, 117– 118 source involvement, 118 –120 North China Craton rocks compared, 120–123 petrography, 103–105 tectonomagmatic significance, 123– 125 multiple vergences, 318– 322 rift basins, 193–196 stratigraphy, 104 structural elements, 318, 321 subduction evidence, 325 tectonic events, 318, 321 Laioning Province Triassic-Jurassic, 202– 204 Jurassic-Cretaceous, 204–205 Cretaceous extension, 205– 207, 214– 216 basins fill and facies, 209–214 drivers, 217– 218 structures, 207–209 timing, 216– 217 tectonic setting, 190– 191, 317– 318 underplating evidence, 325– 326 Yanshan platform-fold belt. See Yanshan orogenic belt Yanshanian orogen, 12, 16, 17, 176, 191, 191 –192 Yellow Sea Bouguer gravity, 283– 284, 286 geophysical survey, 283– 285 Moho depth, 285 seismic velocity, 284, 287

Yinan gabbros, 38 Yingshan– Dingyuan fault, 240 Yinmawanshan granite pluton, 173, 180 Yinshan– Yanshan belt, 240 Yinshan– Yanshan orogen, 181 Yinshan– Yanshan–Liaoxi belt, 79 Yiwulushan granite, 174 Yiwulushan metamorphic core complex, 172 Yixian Formation, 38, 204, 243, 246, 253, 255, 260 Yuanbaoshan arc, 9 Yuantongshan Formation, 243, 248, 249, 253 Yumeng Shan metamorphic core complex, 172 Yunmeng Shan metamorphic core complex, 180, 202, 203 Yunxialing fault, 192 Zhangbei–Guyuan basins, 243 Zhangjiakou Basin, 240 Zhangjiakou fault, 242 Zhangjiakou Formation, 37, 192, 194, 196, 243, 246, 247, 253, 260 Zhangqiao Formation, 243, 251, 252, 253 Zhangwu basalt, 37 Zhaofang pluton, 173 Zhaotun granite pluton, 173, 180 Zhaoyuan– Pingdu shear zone, 176 Zhougongshan Formation, 243, 248, 249, 250, 253, 256 Zhucheng– Rongcheng collision zone, 281 Zhuxiang Formation, 243, 248, 251, 253 Zibo carbonatites, 38 zircon dating, 6, 8 Dandong granite, 175– 180 Inner Mongolian–Daxinganling Orogenic Belt, 227, 228, 229 Jiaodong Peninsula, 132, 312 mantle-lower crustal xenoliths, 58–59, 324– 325 Zuoping gabbro-diorite complex, 37, 38, 39

The North China craton is the only known place where an Archaean craton with a thick tectospheric root lost half of that root in younger tectonism by processes such as delamination, convection, hydration-weakening, compositional change or some other mechanism. In this volume, authors provide data constraining the geometry and timing of root loss, aimed at understanding why and how continental roots are lost in general. Modelling how often this process may have occurred in the geological past, and how much lithospheric material has been recycled to the convecting mantle through this mechanism, could drastically change our current understanding of crustal growth rates and processes. Possible triggering mechanisms for root loss include collision of the South China (Yangtze) and North China cratons in the Triassic, the India–Asia collision, closure of the Solonker and Monhgol–Okhotsk oceans, Mesozoic subduction of the Pacific Plate beneath eastern China, impingement of mantle plumes, mantle hydration from long-term subduction and several rifting events. In this volume, we link studies of crustal tectonics with investigations aimed at determining the nature of and timing of the formation and loss of the root, in order to better-understand mechanisms of continental root formation, evolution and recycling/removal.