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Earth-Science Reviews 65 (2004) 223 每 275 www.elsevier.com/locate/earscirev

A new model for the dynamic evolution of Chinese lithosphere: &continental roots每plume tectonics*
J.F. Deng *, X.X. Mo, H.L. Zhao, Z.X. Wu, Z.H. Luo, S.G. Su
China University of Geosciences, 29 Xueyuan Road, Beijing 100083, PR China Received 23 April 2003; accepted 22 August 2003

Abstract Chinese continental lithosphere comprises three tectonic domains: (1) eastern China, a region characterized by rifting, extensional basins, and voluminous basaltic volcanism; (2) central China, a plexus of cratonic terrains with low-heat flow (40 每 50 mw/m2), including the Tarim, Erdos, and Yangtze blocks welded by pre-Cenozoic orogenic belts; and (3) western China, a region comprising the Qinghai 每 Tibet 每 Himalaya orogen. The relatively thin crust ( f 35 km) and lithosphere ( f 70 km) of eastern China is believed to reflect mantle upwelling, while near-normal crust ( f 45 km) and thickened lithosphere (>200 km) in central China suggest a mantle lithosphere root resembling those inferred for the Kaapvaal and Siberian cratons. Thickened crust ( f 70 km) and lithosphere (>150 km) in western China reflect the advanced development of an orogen root. The lower density and relative buoyancy of central cratonic mantle roots appear to have contributed to the long-term tectonic stability of the central Chinese lithospheric blocks. In contrast, gravitational instability produced by thickening of the denser western Chinese lithosphere and the resulting subsidence and eventual delamination of orogenic roots are believed to have led to postorogenic extensional collapse. Although the geological relationships suggest compressional forces caused both mountain building and orogen root development, extensional stress resulting from gravitational collapse is believed to have induced lithospheric and crustal thinning with concomitant reduction of topography. The term &continental roots 每 plume tectonics* has been adopted to describe the configuration and dynamic condition of subcontinental lithosphere and upper mantle beneath China. Accordingly, it is proposed that supracrustal tectonic forms represent the surface expressions of, and responses to, deep &continental roots 每 plume tectonics*. While the prevailing view is that the western orogenic belt is not genetically related to eastern continental rifting, such a relationship is inherent to the continental roots 每 plume tectonics model. It is further proposed that the formation of orogenic lithosphere roots triggered eastward extrusion of the asthenosphere along the 400-km depth mantle interface, and, in response to subduction at the eastern margin, produces plume-like upwelling beneath eastern China. We suggest that processes involved in continental roots 每 plume tectonics are directly responsible for the formation and evolution of the Eurasian supercontinent. D 2004 Elsevier B.V. All rights reserved.
Keywords: China; Lithosphere; Continental roots 每 plume tectonics

1. Introduction
* Corresponding author. E-mail addresses: dengjf@cugb.edu.cn (J.F. Deng), mxx@cugb.edu.cn (X.X. Mo). 0012-8252/$ - see front matter D 2004 Elsevier B.V. All rights reserved. doi:10.1016/j.earscirev.2003.08.001

Although plate tectonics provide a framework for reconstructing the geological history of Chinese con-

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tinental lithosphere, our understanding of relationships between supracrustal processes and the underlying mantle is still at an early stage. In this review, we explore the results of recent petrological, geological, and geophysical studies in China and propose a conceptual framework for guiding future research on the genesis of continental lithosphere. In general, three main tectonic domain types appear to characterize shallow regions of the earth*s continental lithosphere: compressional orogens, extensional rifted zones, and ancient cratonic shields corresponding to crustal mountain roots and orogenic lithosphere roots, regions exposed to plume-like mantle upwelling, and areas underlain by cratonic mantle lithospheric roots, respectively. In each case, these domains appear to express discrete responses to deeper sublithospheric mantle processes. Coexisting lithospheric domain types in China appear to match those in other continents, suggesting that eastern Eurasia may be a useful analogy for developing generic models for continental crustal genesis.

especially between the upper and lower crust, crust and mantle (Moho), and lithosphere and asthenosphere. From a fluid dynamics standpoint, if discrete systems are constrained by different boundary conditions, the resulting convective patterns will vary. It is therefore important to clarify boundary conditions defining the crust 每 mantle system in a particular region. With respect to Chinese continental lithosphere, the most important geodynamic boundaries appear to be the Siberian and Indian blocks, and the Pacific Ocean. Considered mostly in a two-dimensional (2-D) framework, vertical discontinuities, such as the Moho and lithosphere 每 asthenosphere boundary, are also critical, while the fourth dimension, time, is equally critical. 2.2. Modeling approaches We combine several perspectives in developing a comprehensive dynamic model. The first of these is the notion that erupted magmas and their entrained xenoliths provide thermal and compositional &probes* of their mantle sources. Geochemical evidence from basaltic probes potentially records the changing thermal state of the upper mantle and its collision-related flow behavior, with high-pressure magmas reflecting equilibration conditions in the asthenosphere, and deep crust- and mantle-derived xenoliths reflecting conditions in the lower parts of the lithosphere. Current models, however, are controversial and poorly integrated with 2-D thin viscous sheet tectonic models (e.g., England and Molnar, 1997). Furthermore, mantle signatures are in many cases obscured as a result of contamination by crust 每 mantle mixing during magma transport to the surface. With care, however, the geochemical compositions of magmas may be interpreted as diagnostic of the tectonic, thermal, and compositional attributes of their mantle sources. Moreover, the evidence for P每 T 每t metamorphic paths from exposed Precambrian metamorphic lithologies yields fundamental insights into the interplay of deepseated and shallow responses. A second approach is to integrate the results of the Petrologic probe studies with those from geophysical investigations and plate kinematic reconstructions in developing models for mantle flow fields and their shallow responses. Mantle dynamic models have been

2. Continental roots每 plume tectonics 2.1. Rationale for a new model By reference to P-wave seismic velocity (Vp) or density structure, the &petrologic* structure of continental lithosphere provides a detailed interpretable record of lithosphere tectonic evolution (Wu et al., 1994). Due to their physico 每 chemical differences, diverse lithologies respond differently to similar geodynamic effects, some being more readily reworked through time, others being more resistant. Although petrologic structures mostly record recent tectonic 每 magmatic events, the effects of older events are also preserved, often exposed by denudation. Apparent discrepancies between deep and shallow structural attributes may be considered to reflect processes of heat and mass transfer from the mantle. The effects of these may be complicated, however, given the layered character of crust and upper mantle and varying extent of coupling between layers, in contrast to conditions expected in a homogeneous medium. Structural discrepancies between the upper and lower crust therefore provide a basis for reconstructing and understanding the interactions between different layers,

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mostly developed from geophysical data although interpretations of shallow processes〞tectonic deformation, crustal uplift, sedimentation, basin formation, magmatism, seismicity, and mineralization〞have played an increasing role in constraining crust 每 mantle interactions. For either case, the basis for integrating geological, geophysical, and petrologic data, and testing the validity of models rests on a small number of physico 每 chemical principles that govern phase equilibria, kinetics, and element partitioning. Here, we draw on Cenozoic and Mesozoic examples in developing the &continental roots 每 plume tectonics* model and evaluate these in light of: (1) the assembly of Chinese continental lithosphere during the Mesozoic; (2) the largely complete geological record of recent tectonic events (e.g., the Himalayan orogeny); and (3) the lack of ambiguity suggested by Petrologic-probe, geophysical, and tectonic constraints. 2.3. Continental roots 每plume tectonics 2.3.1. The lithosphere每 asthenosphere boundary A precise definition of &lithosphere base* is handicapped by a lack of agreement on the criteria used for defining the asthenosphere. Here, we adopt the definition of Ringwood (1975) and Condie (1982) that the asthenosphere is a weak, plastic layer extending to c. 700 km depths, divided into upper and lower parts〞

the upper part forming the seismic &low-velocity zone*. In addition to lower P- and S-wave velocities, the latter is characterized by lower seismic attenuation (Q) and higher electrical conductivity, consistent with interstitial (incipient) partial melt, buffered by the dehydration of amphibole and phlogopite (e.g., Lambert and Wyllie, 1970; Condie, 1982). This is consistent with the sharp boundary between the lithosphere and low-velocity zone and higher surface heat flux associated with shallow asthenosphere. Assuming basaltic magma is a partial melt of the mantle, the depth of melt segregation provides a petrologic constraint for the top of the asthenosphere (Deng et al., 1984, 1985). Thus, definitions of the uppermost asthenosphere (hence, base of the lithosphere) are consistent from both petrologic and seismic velocity standpoints. According to Song et al. (1986) (Fig. 1), the lithosphere 每 asthenosphere boundary occurs at 60 每 68 km in depth beneath eastern China (the north China plain, Yellow Sea, and East and South China Seas), c. 118 km beneath the Qinghai 每 Tibet plateau, and c. 104 km beneath its eastern margin, the &North 每 South* tectonic belt. However, there is no low-velocity zone beneath the Yangtze block (i.e., the south China block in Fig. 1). The lithosphere 每 asthenosphere boundary beneath Xingtai (north China) appears at only 83 km in depth (Teng et al., 1982), whereas beneath eastern Tarim, low-velocity mantle is

Fig. 1. Upper mantle Vp structure of China continent (after Song et al., 1986).

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absent down to at least c. 460 km deep (8301 Workgroup, 1988). According to magnetotelluric studies, lithosphere 每 asthenosphere boundary is located at 60 每 80 km in depth beneath the north China plain (Liu, 1985), 100每 130 km between the northern margin of Erdos to Yinshan mountain (Ma et al., 1991), V c. 320 km beneath central Hunan Province (Yangtze block; Rao, 1993), 120每 140 and 160 每200 km beneath the southern and northern parts of the Qinghai 每 Tibet plateau, respectively (Wu et al., 1989), 140 每160 km beneath Qaidam 每Qilianshan 每Beishan (Zhu and Hu, 1995), and 80 km beneath the North 每 South tectonic belt (rising to c. 100 km deep beneath its eastern and western flanks; Ma, 1987). According to the mantle Vs structure, the lithosphere 每 asthenosphere boundary occurs at 60每 80 km in depth beneath eastern China, no low-velocity zone beneath Yangtze block (Song et al., 1992) and Tarim block, 116 每121 and 74 每 92 km beneath Qaidam and Qilianshan, respectively (An et al., 1993; Zhuang et al., 1992), 120每 130 km beneath the Qinghai 每 Tibet plateau and 67 每 74 km beneath the North 每 South tectonic belt (Zhuang et al., 1992), and 90 每 110 km beneath Sanjiang 每 Bayan Har 每 Qaidam (Zhou et al., 1991). Beghoul et al. (1993) suggested that the Qinghai 每 Tibet lithosphere thickness was c. 205每 250 km on the basis of P-wave data by two stations situated within the region. Recently, based on the high resolution surface wave tomography, Zhu et al. (2002, Table 1) presented the average lithosphere thickness as follows: 140每 186 km for Qinghai 每 Tibet, 190 km for Tarim, 160 km for upper Yangtze, 72每 90 km for east China, and 56 每 65 km for marginal seas (South China, Japan, and Okhotsk). Petrologic studies have indicated that Cenozoic basalt magmas were generated mainly at depths c. 60每 80 km beneath the east Chinese rift zone (Deng et al., 1990a), and c. 80 每 130 km beneath the northern margin of the Qinghai 每Tibet plateau (Lai et al., 1996), being consistent in each case with geophysical data and suggesting partial melt segregation from the uppermost asthenosphere. From the above information, we conclude that average lithosphere thickness varies from c. 70 km for eastern China, to c. 200 km for central Chinese cratonic blocks, and to c. 150每 100 km for the Qinghai 每 Tibet collisional belt. Assuming that the lithosphere 每asthenosphere boundary corresponds to Vs 4.30 km/s, marking the depth of alkali basaltic

magma segregation beneath eastern China, the basal lithosphere surface is shown in a series of Vs structure cross-sections (Fig. 2). These indicate thickened lithosphere beneath Tarim, Yangtze, and Qinghai 每 Tibet, referred to as &lithosphere roots*, and thinned lithosphere beneath east China corresponding to upwelling mantle &plumes*. On this basis, &continental roots 每 plume tectonics* provides a generalized framework for describing the Chinese continental lithosphere. 2.3.2. The crust 每 mantle boundary All geophysical data defining the Moho depth beneath China yield a consistent picture, giving average crustal thickness of 30 每35, 40 每 50 km, and c. 68 km for eastern China, central China, and the Qinghai 每 Tibet plateau, respectively, varying gradationally between the first two regions and at margins of the Qinghai 每Tibet plateau (Feng, 1985). Surface elevations correspondingly vary from >4500 m above sea level in the Qinghai 每 Tibet plateau, to 2000 每 1000 m in the area east and north to the plateau, and then to < 500 m of elevation in most of eastern China (Chinese Academy of Geology, 1973). Accordingly, Fig. 3 illustrates the notion that compressional force produces the formation of crustal and lithospheric roots and elevated topography (Fig. 3a); On the other hand, extensional forces produce thinned lithosphere and reduced topography (Fig. 3c). Depicting the analogy of a sailing boat on water, Fig. 3b suggests that buoyancy of the craton keeps it afloat in the denser asthenosphere. Because surface tectonic forms appear to support this type of mechanistic model, it seems reasonable to conclude that continental supracrustal domains represent surface expressions of, and responses to, the effects of deep, &continental roots 每 plume* tectonics. 2.3.3. Cenozoic lithosphere domains It is widely accepted that Chinese continent is divided into western (&Himalayan*) and eastern (&circum-Pacific*) sectors bounded by a north 每 south tectonic belt at c. 102 每 105jE (Huang et al., 1980; Ma et al., 1987; Wang, 1982). However, this paper proposed the following division of three Cenozoic geotectonic units or domains of Chinese continental lithosphere (Fig. 4): rift zone of eastern China, Cratonic blocks of central China, and collisional orogenic belt of western China, as predicted by the continental root 每plume

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Fig. 2. Lithospheric base contour of China continent. Pictured by the upper mantle Vs structure (data from Song et al., 1991, 1992; An et al., 1993). A〞asthenosphere; L〞lithosphere; LV〞residual fragment of lithosphere.

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Fig. 3. Cartoon of China lithosphere/crust contour. (a) The orogenic belt of Qinghai 每 Tibet. (b) Cratonic block in central China. (c) Continental rift in east China. S〞surface; M〞Moho; L/A〞the boundary between lithosphere and asthenosphere.

tectonics model (Deng et al., 1996a, 1997a). The western boundary of stretched (thinned) lithosphere extends northwards from Beihai (Guangxi Province)

through Sanshui (Guangdong), Mingxi (Fujian), Quxian (Zhejiang), Nanjing (Jiangsu), Feixiang (Hebei), Jinning (Inner Mongolia), and to Mongolia and Lake

Fig. 4. Geotectonic units of China on a lithosphere scale since Cenozoic (after Deng et al., 1996a). 1〞The boundary of volcanic rocks and geotectonic unit; I〞east China continental rift zone; II〞central China blocks; III〞Qinghai 每 Tibet orogenic belt. 2〞Upper mantle isodensity layer, density boundary (after Wang and Cheng, 1982); layer density: r = + 1, layer thickness: 150 每 850 m; layer density: r = ? 1, layer thickness: 150 每 >850 m; layer density: jrj = 1, layer thickness: 9 每 150 m.

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Baikal (Fig. 4). To the east, thinned eastern Chinese lithosphere is characterized by Cenozoic alkali basalts with abundant mantle-derived xenoliths (Deng et al., 1984, 1985). The shoshonite volcanism, distributed along Kunlun, Qilian, and Sanjiang mountains, marks northern and eastern margins of the Qinghai 每Tibet plateau, associated with lithosphere-scale compressional tectonics (Fig. 4). There is little or no Cenozoic volcanism in tectonically stable central Chinese units such as the Tarim, Alxa, Erdos, and Yangtze blocks, separating the eastern and western volcanic provinces. Although considerably smaller in size than most cratons in the world, their long-term tectonic stability (Fig. 4), as well as lack of recent seismicity (except at their margins), integrally uplifting, and low-surface heat flow (40 每

50 mw/m2), is clearly shown from the geological record. Moreover, isostatic gravity anomalies reflect crustal and mantle heterogeneity, indicating long- and short-wave length density variations, respectively. While 1 ? 1j free-air gravity anomaly contours correspond to topographic highs and lows, 3 ? 3, 5 ? 5, 7 ? 7, and 9 ? 9j free-air anomalies are more regular and indicate deep-seated mantle density anomalies (Ding, 1991). There are three gravity anomaly regions in China, i.e., a positive east 每westtrending anomaly associated with the Qinghai 每 Tibet plateau, a negative anomaly extending from Xinjiang through Erdos to Sichuan 每 Yunnan, and an arcshaped, eastward convex positive anomaly in eastern China (Wang and Cheng, 1982; Ding, 1991). These anomalies are believed to reflect dynamic nonequilib-

Fig. 5. The P 每 T curves of granulite and eclogite facies. Facies transitional boundary: (a) between amphibolite and middle-pressure granulite facies; (b) between middle- and high-pressure granulite facies; (c) between granulite and eclogite facies of basalt rocks; (d) between granulite and eclogite facies of granite rocks. (a) and (b) and (c) from Holloway and Wood (1988); (d) from Deng (1987) and Wyllie (1981). The dashed lines are geotherm (cf. Deng et al., 1995a). 1-Qaidam; 2-Northern margin of Qaidam; 3-Qilianshan; 4-eastern Jiuquan Basin; 5-Huahai Basin; 6-Beishan.

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rium associated with recent redistribution of mantle mass (Wang and Cheng, 1982; Ding, 1991). For example, the Qinghai 每Tibet and eastern China positive anomalies are considered to reflect the respective motions of Indian and Pacific plates in relation to Eurasia (Wang and Cheng, 1982). Using the 5 ? 5j anomalies, Wang and Cheng (1982) mapped the thickness of the upper mantle isodensity layer, indicating two regions of higher mantle density beneath

Table 1 Cenozoic lithosphere-scale tectonic units of China (after Deng et al., 1996a) Qinghai 每 Tibet intracontinental orogenic belt Crustal thickness (km) Lithosphere thickness (km) Upper mantle density (g/cm3) Elevation (m) 70 每 80 (crustal root) 150 (orogenic lithosphere root) 3.4 每 3.65 Central China cratonic group 40 每 50 (normal) 200 (cratonic lithosphere root) 3.10 每 3.22 East China continental rift zone 30 每 35 (thinned) 70 (mantle plume)

3.23 每 3.30

Magmatism

Surface heat flow (mw/m2)

Earthquake

>4500 (plateau, high mountain) Intracontinental shoshonitic series, Ms/two mica granites Variable; north part: 40 每 50; margins: 70 每 90; inner part: 100 每 300 Many and strong

2000 每 1000 (basin and plateau) None

V 500 (basin and range) Continental rifting basalts

40 每 50

60 每 70

(1) the Qinghai 每 Tibet plateau west to Xining 每 Chendu每 Kunming and (2) the eastern China east to Changchun 每 Taiyuan 每 Guangzhou, separated by a lower density region beneath Xinjiang 每 Erdos 每 Sichuan of central China. The upper mantle density distribution from the Moho down to 120 km in depth based on a threedimensional gravity inversion derived from the l ? 1j average Bouguer gravity anomaly (Feng, 1985) indicates three upper mantle density regions which appear to be accordant to the overlying crustal block structure. These are 3.40 每 3.65 g/cm3 (Qinghai 每 Tibet plateau), 3.23 每 3.30 g/cm 3 (eastern China), and 3.10 每 3.22 g/cm3 (central China〞corresponding to the Sichuan basin, Erdos plateau, Shanxi graben, Tarim basin, and Junggar basin). The long-wave isostatic gravity anomaly suggests that upper mantle density is consistent with our lithosphere-scale Cenozoic geotectonic scheme. Accordingly, we adopted boundaries between higher density (layer density r= + 1), normal density (ArA = 1), and lower density (r = ? 1) upper mantle regions from Wang and Cheng (1982, Fig. 5) in Fig. 4, each of which corresponds roughly with the volcanic boundaries noted above. This scenario is consistent with the cratonic buoyancy model, reflecting significantly lower mantle densities beneath central Chinese cratonic blocks compared to the Qinghai Tibet plateau, inferred from the lithosphere/crust contour (Fig. 3b). Thus, the density character of cratonic (central Chinese) and orogenic (Qinghai 每Tibet) lithosphere roots appears to be the prime determinant of their respective gravitational stability and susceptibility to delamination, respectively. These features are summarized in Table 1.

Recent tectonics

Transpressional

Low-velocity and highconductivity layer in crust

Present

None (except for interblock boundaries) Stable (whole uplifting and subsiding) Absent

Many and strong

3. Chinese lithosphere structure Petrologic approaches to studying crust 每 mantle structure involve at least three types of investigations (Deng et al., 1992a) of: (1) surface or near-surface exposures of deep, mainly Precambrian, metamorphic crustal rocks; (2) basalt-hosted crustal and mantlederived xenoliths; and (3) xenolith- and lava-derived thermobarometric and compositional information (Wu et al., 1994). For regions affected by multiple magmatic 每 tectonic events, crust 每 mantle relationships

Transtensional

Present

J.F. Deng et al. / Earth-Science Reviews 65 (2004) 223每275 Table 2 Vp and Vs of minerals (km/s) Quartz Vp Vs 6.05 4.09 Alkali feldspar 6.01 3.34 Plagioclase (An = 24) 6.22 3.40 Plagioclase (An = 53) 6.57 3.53 Biotite 5.26 2.87 Amphibole 7.04 3.81 Diopside 7.70 4.38 Garnet 8.53 4.76

231

Olivine 8.32 4.57

Note: Ol after Birch, 1961; others after Christernsen and Fountain, 1975.

have been reconstructed using both petrologic and geotectonic approaches (Deng et al., 1992a; Wu et al., 1994). 3.1. Continental crust According to these studies, three types of crustal structure appear to characterize China, as represented by eastern China, central China, and the Qinghai 每 Tibet 每 Himalaya plateau (Table 5), consistent with conventional models assuming here to comprise greenschist, amphibolite, and granulite facies lithologies, respectively (Fountain and Salisburg, 1981; Wu and Guo, 1991). However, Qinghai 每 Tibet 每Himalayan crust possesses a four-layered structure, and the fourth layer (&thickened lower crust*) is referred to a high-pressure granulite F eclogite facies &mountain root* (Deng et al., 1995a, 1997a). The four-layered Vp structure of &double* thickness of Qinghai 每Tibet orogenic crust represents an extreme case with which to compare &normal* crustal structure. As noted, experimental and thermodynamic data suggest that critical metamorphic reactions occur with increasing pressure during the transition from hypersthene-bearing granulite to eclogite (Deng, 1987; Holloway and Wood, 1988). At 400每 800 jC, 1.0每 1.1 GPa pressure (c. 35每 40 km depth), high-pressure

granulite forms in response to orthopyroxene breakdown, while at 400 每 800 jC, 1.2 每 1.6 GPa pressure (c. 45 每 60 km depth), eclogite forms in response to plagioclase breakdown (Fig. 5). Thus, thickened lower crust may comprise of high-pressure granulite F eclogite facies rocks. Due to the higher Vp and lower Vs values of feldspar compared to quartz (Table 2), feldspar- and quartz-rich crust may be distinguished from crustal Vp/Vs profiles. Accordingly, thickened lower crust probably comprises of highpressure granulite of intermediate to granodioritic composition, whereas mountain roots are more likely to consist of &granitic* eclogite facies assemblages (Tables 2 每4). Seismic profiles also indicate that the petrologic structure of &double* Himalayan crust (Deng et al., 1995a) resembles that of the Qilian orogen. In contrast, cratonic lower crust has a high-velocity layer with Vp up to 6.8 km/s, suggesting high-pressure granulite of granodioritic composition. Lowermost crust in the continental rifting zone comprises a high-velocity layer (Vp of 7.2 每7.4 km/s), interpreted to represent mantle-derived basaltic underplating (Fountain, 1989; Wu et al., 1994) similar to the formation of the lowermost crust in Gangdese and Bayan Har of the Qinghai 每 Tibet Plateau (Cui, 1987; Deng et al., 1997a,b).

Table 3 Whole-rock Vp and Vs calculated from mineral data of Table 2 (after Deng et al., 1996a) Rock Basic granulite (Opx 20; Cpx 20; Ga 10; Pl 50) 7.25 4.01 Intermediate granulite (Opx 15; Cpx 15; Ga 5; Pl 65) Intermediate granulite (Opx 10; Cpx 10; Bi 10; Ga 5; Pl 65) 6.55 3.65 Acid granulite (Opx 2; Cpx 3; Ga 5; Pl 60; Qz 30) 6.36 3.73 HP intermediate granulite (Cpx 25; Ga 10; Pl 55; Qz 10) Basic eclogite (Ga 50; Cpx 50) Acid eclogite (Cpx 15; Ga 15; Qz 70) Harzburgite (Opx 20; Ol 80) Lherzolite (Cpx 10; Opx 20; Ol 70)

Vp (km/s) Vs (km/s)

6.80 3.82

6.80 3.85

8.12 4.57

6.67 4.23

8.23 4.61

8.16 4.59

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Table 4 Experimental Vp and Vs of natural rocks at 0.6 GPa Rock Vp (km/s) Vs (km/s) Granite 6.05 每 6.50 3.59 每 3.80 Amphibolite 每 facies gneiss 6.30 每 6.60 3.75 Amphibolite 6.92 每 7.18 3.90 Intermediate granulite 6.70 每 6.80 3.90 Basic granulite 7.10 每 7.40 3.90 每 3.93 Acid granulite 6.41 每 6.57 3.58 每 3.74 Basic eclogite 7.90 每 8.20 4.44 每 4.46 Dunite 8.10 每 8.45 4.58 每 4.83

Note: Data from Wu and Deng (1994), Fountain (1976), Kern (1982), Christernsen and Fountain (1975), Birch (1961).

3.2. Upper mantle The combined evidence of xenoliths and seismic profiles indicates that tectonic domains are also distinguished by their lithospheric mantle compositions, i.e., spinel lherzolite, with trapped melt products such as pyroxenite beneath thinned, extensional crust (Table 5), refractory harzburgite beneath cratonic blocks (Wu et al., 1994), and garnet peridotite with eclogite beneath thickened orogenic regions (Deng et al., 1995, 1997a). Low Vs velocity character often encountered at about 80 km depth in the uppermost subcratonic mantle is attributed on the basis of theoretical calculations (DVp of + 0.199 km/s and DVs of ? 0.227 km/s) to the spinel每 garnet transition (Wu et al., 1994; Deng et al., 1997a,b). The asthenosphere beneath both extensional rift zones and the Qinghai 每 Tibet orogenic belt is believed to consist of garnet peridotite with interstitial basaltic melt. However, the general absence of volcanism in cratonic blocks suggests that asthenospheric melt is not trapped. 3.3. Lower crust and Moho formation Although the overall distinction between thinned extensional crust and thickened orogenic crust is clear (Table 5), seismic profiles suggest that their upper and middle layers have not changed significantly, thickening and thinning being confined mostly to the lower crust. The transition from brittle to ductile deformation in granitic crust occurs between c. 500 jC (at 0.4 GPa pressure) and 400 jC (at 0.3 GPa pressure) (Wu and Deng, 1994), approximating the 500 jC crustal isotherm at c. 20 每25 km depth (Wu and Deng, 1994). Thus, the Moho beneath the regions of normal and

thinned crust may be regarded as a compositional interface, whereas beneath thickened orogenic belts, the Moho probably represents a physical (rheologic) boundary. During orogenic crustal thickening and mountain root formation, mafic crustal lithologies buried in response to compression and hydrostatic stress would reequilibrate as eclogite (Fig. 5). In contrast to granitic compositions, mafic eclogite shows brittle deformation at lower crustal pressures. Therefore, during lower crustal thickening, eclogitic components may be drawn downward as fragments enclosed by plastic lower crustal matrix (Fig. 6a). When compressional forces decrease or cease, eclogite fragments may accumulate above the preexisting Moho on a peridotite base as a consequence of the density difference between basic and granitic &eclogite* lithologies or high-pressure granulite. Because of much larger difference in density and seismic velocity between basic and granitic &eclogite* than that between basic eclogitic cumulates and peridotite, the sharp break in density and seismic velocity migrate to the interface between basic eclogitic cumulates and granitic &eclogite* or high-pressure granulite, producing a &new* Moho in place of that defined by a peridotite base (Fig. 6b)〞an effect which may be termed &dynamic-gravitational* differentiation (Deng et al., 1995). During crustal thinning associated with processes of orogenic collapse and cratonization, basic eclogite cover of the upper mantle and granitic &eclogite* in the mountain root may reequilibrate into granulite facies, resulting in yet another Moho〞an expression of changing metamorphic facies conditions during a given magmatic 每tectonic event. In the case of extremely rapid uplift following mountain root forma-

Note to Table 5: Notes: U: upper; M: middle; L: lower; T.L: thickened lower crust; M.R.: mountain root; L/A: boundary between lithosphere and asthenosphere; Chl.F.: greenschist facies; Am.F.: amphibolite facies; Py.F.: granulite facies; HP Py.F.: high-pressure granulite facies; g: granitic; y: dioritic; b: basaltic; gy: granodioritic; LVZ: low-velocity zone.

J.F. Deng et al. / Earth-Science Reviews 65 (2004) 223每275 Table 5 Crust 每 mantle petrological structure of Chinese lithosphere domains (after Deng et al., 1996a)

233

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Fig. 6. Differentiation of deep crustal materials (a) and formation of new Moho (b). 1〞Middle-crust base; 2〞preexisting Moho; 3〞 basic eclogite; 4〞granitic eclogite and high-pressure granulite; 5〞 peridotite; 6〞new Moho.

tion, eclogite fragments in the latter and the subjacent upper mantle may not completely reequilibrate into granulite surviving, if exposed, as shallow crust. Thus, both the mountain root formation and ensuing rapid uplift may be regarded as a possible mechanism for the formation and transport to the near surface of ultrahigh-pressure (UHP) metamorphic rocks. 3.4. Tectonic implications It is widely accepted that primitive basalt represents the minimum-melt composition of upper mantle peridotite as expressed by much of the oceanic crust, immature volcanic arcs, and underplating of continental crust. Granites likewise resemble minimum-melt compositions within the basaltic system as expressed in mature volcanic arcs and continental crust, in many cases, at least, produced by partial melting of basaltic crust. Because of the greater buoyancy of granitic and granodioritic crust compared to juvenile basaltic crust, continental lithologies are preferentially stabilized on the earth*s surface. Likewise, harzburgite residues

generated by the removal of basaltic melt from primitive lherzolitic mantle are characterized by increased Vp and Vs values (Tables 2 and 3) and decreased density due to their higher Mg/(Mg + SFe) values. The low crustal density and lack of a crustal lowvelocity layer together with thickened low-density subcrustal lithosphere (Table 5) all contribute to this boyuancy effect, being consistent with gravity models (Table 5, Fig. 4). In contrast, higher-density granodioritic crust, juvenile subcrustal basaltic underplating, density-overturn derived from the crustal low-velocity zone, shallow asthenosphere, and density-overturn resulting from asthenospheric interstitial melt all contribute to the relative tectonic instability exhibited by thinned extensional eastern Chinese lithosphere. The &basification* of eastern Chinese crust has been attributed to Mesozoic and Cenozoic tectonic reactivation (Wu et al., 1994). In contrast to central Chinese cratonic blocks, mantle lithosphere beneath the Qinghai 每Tibet orogenic belt (garnet lherzolite + eclogite) is significantly denser and produced, presumably as a consequence of crustal thickening, depression of the Moho and extensional lithospheric detachment. These effects would, moreover, provide an opportunity for the recycling of mountain root-derived crustal materials, in turn, adding a felsic component to the mountain root contributing rapid uplift inferred for the Himalayan and Qilian orogens. The widespread occurrence of granitic mylonite in the middle crust (Deng et al., 1995a) is regarded to be a record of intensive intracrustal tectonic activity (Table 5).

4. The record of magmatic activity 4.1. Igneous rock relationships The correspondence of igneous rock type and tectonic setting is well established. However, despite numerous studies of rift- and plume-related (e.g., Condie, 1982; Morgan and Baker, 1983), and kimberlite (e.g., Boyd and Gurney, 1986) volcanism, there are relatively few synthetic studies of intracontinental magmatic activity. As a contribution to this review, we have summarized the apparent relationships among magma composition, intracontinental domain type, and the continental roots 每plume hypothesis (Table 6). We note that lithologic assemblages in &welded*

J.F. Deng et al. / Earth-Science Reviews 65 (2004) 223每275 Table 6 Intracontinental igneous petrotectonic assemblages (after Deng et al., 1996a) Tectonic setting Petrotectonic assemblage Roots 每 plume tectonics type Orogenic lithosphere root and mountain root

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Intracontinental (1) Oceanic orogenic belt closure and continent 每 continent collision orogeny: blueschist and eclogite; collisiontype PTt paths; calc 每 alkaline and high-K calc 每 alkaline magmatism; (2) Intracontinental orogeny: intracontinental subduction boundary: biotite/two mica/ muscovite granites; distributed shortening 每 thickening boundary: shoshonitic series magmatism (3) Orogenic collapse: late or postorogenic A-type granites Early stage: alkaline syenite and quartz syenite (no negative Eu anomaly), peralkaline graniteLast stage: alkaline syenite and quartz syenite (negative Eu anomaly), peralkaline granite Continental Alkaline olivine basaltic rifting zone series, continental flood basalts, nonorogenic A-type granites Craton Kimberlite or no magmatism

De-rooting: (1) Early stage: lithosphere thinning with crustal root (2) Late stage: lithosphere thinning without crustal root

erably younger than most continental assemblages〞 oldest ocean basement being less than 200 Ma〞it is assumed that the bulk of oceanic lithosphere produced has been consumed by subduction. While oceanic spreading presumably started before a2, subduction had already started since a2 (Table 7). However, the younger (O3) ages of exposed blueschist terranes suggest that only latest-stage subduction products are preserved in intracontinental domain boundaries. Source lithologies for high-pressure metamorphic rocks reflect a range of typically oceanic (backarc) and volcanic arc-related magmatic products. Preservation of the latter probably implies their suprasubduction origins, with blueschists forming under temperatures between c. 200 and 400 jC. For example, Linzizong volcanic and related granitic rocks located at the southern margin of Gangdese are considered to be collision-related products of oceanic closing. Paleocene and Eocene volcanic (64 每 41 Ma) and granitic (55 每 41 Ma) rocks (Liu et al., 1990) range from calc 每 alkaline to high potassic types (Lai et al., 1996). Orogenies invariably follow the termination of subduction rollback and collapse (or consumption by subduction) of backarc basins, inevitably preserving some aspects of magma production and supply systems. 4.3. Muscovite/two-mica granites Miocene muscovite/two-mica granites distributed across 2000 km of the Himalayan orogen (Le Fort, 1981) are regarded as products of collision-related crustal melting. Debate has focused on whether melting occurred at the base of thickened &double* crust or resulted from intracontinental subduction. Assuming that &muscovite/two-mica*-type granitic magma is generated by partial melting of suprasubduction pelitic rocks, the petrologic character of these granites〞 such as interstitial muscovite, widespread perthite development, chemical and mineralogic homogeneity, and the occurrence of wollastonite in outer contact aureoles〞indicates H2O-undersaturated melts at temperatures up to 750 每800 jC rather than the lowtemperature H2O-saturated melts (cf. 600每 650 jC). The characteristics of intruded granitic magmas〞 such as high SiO2 content, initial 87Sr/86Sr ratios and V 18O ratios, K20 z Na2O, and A1203-oversaturation〞and lack of accompanying emplacement of

Mantle plume

Continental lithosphere root

continent 每continent collision zones are invariably of oceanic affinity dominated by ophiolites. 4.2. Continental collision assemblages Eclogitic blueschists exposed in the early Paleozoic Qilian orogen are believed to represent magmatic products generated during oceanic closure immediately prior to a continent 每 continent collision (Table 7; Deng et al., 1996b). Because oceanic crust is consid-

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Table 7 Sequence of events: north Qilian Caledonian orogenic belt (after Deng et al., 1996a) Seafloor spreading and subduction (a) MORB, Pt2, O1, O2, O3 Oceanic closing and continent 每 continent collision (a) Blueschist (eclogite) (subducted oceanic crust; arc volcanic rocks, pelitic rocks), 460 每 440 Ma (O3) (b) Marine molasse, S1, unconformable contact with pre-Silurian Intracontinental subduction orogeny Two mica-granites, 417 每 404Ma(S2 每 S3) Postorogenic collapse and uplift (a) A-type granites, D1, D2 (b) Retro-metamorphism of blueschist, 400 每 380Ma (D1 每 D2) (c) Continental molasse, D1, conformable contact with pre-Devonian Cratonization

Young platform, C

(b) OIB, Pt2, a2, O1, O2, O3

(c) Subduction arc volcanism, a2, O1, O2, O3 (d) Subduction arc granites, 532 每 443 Ma (a2 每 O3)

intermediate or basic magma support their provenance by partial melting of pelitic crustal rocks without the involvement of mantle materials. Intracontinental subduction is a possible mechanism contributing to crust recycling (Deng et al., 1994a,b). Strong negative Eu anomalies in chondrite-normalized REE patterns of muscovite/two-mica granites suggest that these melts equilibrated with plagioclase-bearing assemblages. This, in turn, indicates that the magma was probably not generated in thickened continental crust because plagioclase is

probably absent at depths>50每 60 km, and Himalayan crust having been 60每 75 km in thickness. Experimental petrologic data show that muscovite-bearing melts are generated by partial melting of pelitic compositions at depths of 20 每40 km. The Himalayan muscovite/two-mica granites were formed in 7每 28 Ma, whereas retrograde metamorphism of the central crystalline core occurred in 10 每 20 Ma. According to the results from the deep seismic reflection INDEPTH program (Zhao and Nelson, 1993), the Main Central Thrust (MCT) is likely the locus of intracontinental

Fig. 7. Temperature distribution of higher-Himalaya (after Deng et al., 1994a,b). Dashed line is geotherm of heat-flow = 90 mW/m2. The dashed frame is the temperature and pressure range of low-amphibolite facies or Ky zone. Solid short-line is T and P range of high-amphibolite facies or Sil zone of this area. The solid frame represents T and P range of migmatization. Curve 1 is solidus of granite-excess H2O system. Curve 2 is solidus of muscovite-dehydration. Curve 3 is liquidus of muscovite granite melt with 10% H2O.

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subduction. Therefore, the foreland fold zone, including the Lesser Himalaya, MCT, the higher Himalayan central crystalline core and two-mica granite belt, most likely represents the orogenic expression of intracontinental subduction, comprising a subducted continental block, an intracontinental subduction zone, and an overlying continental block, resepectively (cf, Fig. 8). The two-mica granite belt thus represents the petrologic response to intracontinental subduction. Relevant phase equilibrium data allow prediction of melting conditions in relation to the high Himalayan geotherm (Fig. 7).The implied subduction-related thermal structure in relation to dehydration reactions and muscovite/two-mica granite formation is shown in Fig. 8. The model assumes: (1) an approximately 5km-thick sedimentary cover on the subducted slab; (2) a subduction angle of about 20j; and (3) a slab sedimentary cover continuing down to 30 每 35 km depth. Because the sedimentary cover with lower density could not penetrate the Moho, the temperature rise must have induced dehydration, and melting of pelitic sediments must have resulted from subduction. Initially, water released from kaolinite dehydration reached shallow levels near the MCT, giving rise to the widespread hot springs observed in this region. Water released from both kaolinite and (at deeper levels) pyrophyllite dehydration ascended to the lower part of the overlying plate, triggering H2O-saturated granite formation by transgressing the H2O-saturated granite solidus as shown by curve 3-3 in Fig. 8. As a

result of their low temperature, the granitic melts tended to be trapped at depth, forming large &in situ* plutons and migmatites. In the absence of excess water, partial melting of subducted sediment would clearly not occur at the H2O-saturated solidus. However, continued subduction would result in dehydration of muscovite, as shown by curve 4-4 in Fig. 8, triggering formation of H2O-undersaturated muscovite/two-mica-type granite at temperatures of 750 每 800 jC. By this stage, the bulk of granitic melt would presumably be removed by filter pressing and emplaced into the continental crust. 4.4. Orogenic shoshonites Cenozoic volcanism at localities on the northern margin of the Qinghai 每Tibet plateau, including Kunlun, Hoh Xil, and Yumen, belong to the shoshonitic series, comprising shoshonite, latite, trachyte, and rhyolite (Lai et al., 1996; Turner et al., 1996; Flower et al., 1998). Although these rooks show similar SiO2alkali character to &subalkaline* and alkali olivine basalt series (Fig. 9a), lack of Fe enrichment trends (Fig. 9b) and often high-potash characteristics in them (Fig. 9c and d) confirm their shoshonitic and highpotash calc 每 alkaline affinity. Despite petrologic and geochemical indications of volcanic rocks with island arc 每 continental margin affinity, their emplacement over 1000 km north of the MTB c. 20 每30 Ma, after the &hard* collision of India and Asia, suggest that they represent a distinct

Fig. 8. Thermal structure and magma generation of the intracontinental subduction zone (after Deng et al., 1994a,b). Solid curve is isotherm. Dot 每 dashed curve: 1-1〞kaolinite dehydration reaction; 2-2〞pyrophyllite dehydration reaction which represents the beginning of greenschist facies; 3-3〞granite solidus at excess H2O; 4-4〞granite solidus by muscovite dehydration; 5〞migmatitite and muscovite granite intrusion in and semi-in-situ; 6〞muscovite granite intrusion and magma source; 7〞the water released from dehydration.

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Fig. 9. Chemistry of Cenozoic volcanic rocks of intracontinental orogenic shoshonite series, Qinghai 每 Tibet plateau; I〞high potassic; II〞 potassic; III〞sodium.

&intracontinenal orogenic shoshonite* series. Comparing &high-field strength element* (HFSE; e.g., TiO2, Zr, Nb, etc.) contents of shoshonites from the northern Qinghai 每 Tibet plateau margin with those from typical volcanic arcs, it shows that they both share characteristics of relative depletions in these elements, although HFSE contents in the former are generally higher and transitional to continental rift-type compositions. Although intracontinental orogenic shoshonites share both arc and intraplate geochemical character, clinopyroxenes in Qinghai 每 Tibet shoshonites are also relatively rich in TiO2 (1.04 每 1.74%) compared to those in arc lavas, resembling clinopyroxenes in rift-

related basalts. Thus, compared to oceanic subduction-related arc volcanics, intracontinental orogenic magmas: (1) are mainly shoshonitic rather than calc 每alkaline; (2) are characterized by relatively high whole-rock HFSE contents; and (3) show Ti-rich clinopyroxene compositions. Qinghai 每 Tibet orogenic shoshonites are mostly evolved as indicated by their relatively low Mgnumbers [Mg/(Mg + Fe+ 2)]. However, primitive tephrites with Mg-numbers up to 0.65 probably represent near-primary asthenosphere-derived magmas segregating from the depths of c. 80 每130 km (Lai et al., 1996), being consistent with geophysical constraints

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Fig. 10. Cartoon showing orogenic lithospheric delamination and magma underplating (after Deng et al., 1995a).

on asthenosphere depth of c. 90每 120 km (An et al., 1993) or c. 140 每160 km (Zhu et al., 2002). In contrast to muscovite/two-mica granites, most Qinghai 每Tibet orogenic shoshonitic rocks, generated in the lower parts of thickened continental crust (>50 每 60 km depth), have no negative Eu anomalies. Average crustal thickness in the northern Qinghai 每 Tibet orogenic belt is 60每 70 km with a high-velocity layer (Vp = 7.3 每7.4 km/s) at the base (Cui, 1987), being consistent with basaltic underplating. A model for magma generation in Yumen, being consistent with the relevant petrologic, geological, and geophysical data, is shown in Fig. 10 (Deng et al., 1995a). During lithosphere convergence, Beishan was a part of the Tarim 每 Alxa stable craton with more stable and stronger lithosphere than that of the Qilian orogen. Therefore, at the boundary between Qilian Shan and Beishan, horizontal shortening presumably drove down Qilian rather than Beishan mantle lithosphere. Extensional fracturing along the zone of maximum flexing probably led to the delamination of denser orogenic lithosphere and triggered asthenosphere upwelling. Decompression of the asthenosphere beneath such weak zones separating orogenic and cratonic lithosphere clearly favors partial melting as reflected by underplating at the base of the crust and basalt volcanism at Yumen. 4.5. Postorogenic A-type granites Recent studies by Hong et al. (1991), Rogers and Greenberg (1990) and Eby (1992) suggest that A-type

granite may occur at late- and/or post-postorogenic stages. The advent of lithosphere extension may be regarded as a terminal phase of compressional deformation cycles, whereby the crust and lithosphere become thinned as collapse progresses (e.g., Turner et al., 1992). A-type granites developed in orogenic settings probably signify such a collapse stage, marking continental lithosphere consolidation and tectonic stabilization. Based on the model of continental roots 每 plume tectonics, processes leading to two groups of the postorogenic A-type granites are clearly initiated in the mantle but progress to the crust as extensional stresses become dominant. For example, the Gangdese part of the Qinghai 每 Tibet 每 Himalaya orogen recently entered a stage of orogenic collapse, reducing the lithosphere thickness to c. 120 km from c. 150 每 200 km in its northern and southern parts but keeping the mountain root of c. 80 km. Thus, the progression from orogenic collapse via elimination of mountain roots to the generation of two groups of postorogenic A-type granitic melts is a critical aspect of the continental roots 每plume tectonics model (see Table 6). A-type granitoids include the alkaline syenite 每 quartz syenite series and also peralkaline granites, and their volcanic equivalents include alkaline trachyte 每 quartz trachytes, and peralkaline rhyolites. Experimental-phase equilibrium studies (Huang and Wyllie, 1981, 1986; Stern and Wyllie, 1981; Deng, 1987; Deng et al., 1998a) suggest that: (1) granites and syenites (sensu stricto) form by partial melting at pressures of V 1.0 GPa (in 30 每 40-km-thick crust) and c. 1.5 GPa (in z 55-km-thick crust), respectively;

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(2) plagioclase is absent from the solidus at c. 1.7 GPa (in c. 60-km-thick crust) and absent at subsolidus temperature at 1.5 GPa (having reacted to pyroxene); (3) andesite liquidus phases are P1 F Py at V 1.0 GPa and Py F Ga F Hb at >c. 1.0 GPa, respectively, whereas liquidus plagioclase in basaltic magma occurs at lower pressures ( V 0.1 每0.2 GPa; Ringwood, 1975). The phase equilibria thus indicate that granitic (including peralkaline) melts are generated in the lower crust (if the crust is &normal* in thickness or &thinned*) or middle or upper crust (if the crust is thickened in orogens) and equilibrated with plagioclase during either differentiation or partial melting (indicated by pronounced negative Eu anomalies). Likewise, alkaline syenite 每 quartz syenite without negative Eu anomalies can be interpreted to have formed within the mountain root in the absence of plagioclase, whereas syenites (sensu stricto) which show small negative Eu anomalies formed at depths of about 55 km (c. 1.5 GPa pressure) due to small amounts of plagioclase in the crustal residue, but, at shallow depths, syenite magma formed by differentiation of basaltic or andesitic magma would have negative Eu anomalies due to the presence of liquidus plagioclase between c. 1.0 and 0.1 GPa pressure. Thus, the presence or absence of negative Eu anomalies in alkaline or quartz syenites allows recognition of whether their crustal sources were within the mountain root or not, implying, in turn, that Eu anomaly-free A-type syenites reflect the early stages of orogenic collapse, whereas those with strong negative Eu anomalies indicate later stages of collapse or incipient continental rifting (Table 6). Distinction of the last two conditions, late-stage collapse and continental rifting, requires additional criteria relating to geological setting and the nature of coexisting basalts (Eby, 1992; Rogers and Greenberg, 1990). The presence or absence of Eu anomalies is also suitable for distinguishing other intermediate-acid igneous rocks including andesite, trachyandesite, latite, dacite, quartz trachyte, trachyte, and their intrusive equivalents. These magmas are generated either directly from crustal rocks or via the interaction of mantle-derived basaltic melts and crustal rocks. At depths >c. 50 每 60 km, the crust will conform to plagioclase-free eclogite facies (Fig. 5); therefore, such melts generated from mountain root sources would still lack Eu anomalies. Thus, intermediate-

acid igneous rocks lacking negative Eu anomalies would also signify sources in an orogen-related mountain root. Those intermediate-acid magmas showing Eu anomalies would have formed in the crust with normal thickness or the middle to upper parts of &double* thickness orogenic crust. Thus, the presence or absence of negative Eu anomalies provides a simple way to distinguish orogenic sources with/without mountain roots, and thickened/thinned crust provides, in combination with geochemical, petrologic, and geological data, meaningful constraints on the continental roots 每 plume tectonics model and continental dynamic processes in general.

5. Orogen dynamics of the Qinghai 每 Tibet每Himalaya 5.1. Paired igneous belts Following the India 每Asia &hard* collision (c. 45 Ma), the Qinghai 每 Tibet 每Himalaya orogen entered a new &intracontinental* stage. A notable characteristic of magmatic activity in the region is that igneous rocks are distributed only on the orogen margins, the Yutian每 Yumen volcanics on the northern margin of the Qinghai 每 Tibet plateau and higher Himalaya muscovite/two-mica granites to the south, whereas the broad interior remains relatively free of igneous activity. Thus, distribution of intracontinental igneous activity may be used to define the temporal每 spatial extent of orogen boundaries. Another obvious feature is the apparent pairing of discrete igneous rock belts (Fig. 11; Deng et al., 1994d, 1996a). During the Oligocene (E3), the volcanic zone marking the northern Qiangtang margin appears to be paired with a muscovite/two-mica granite zone on the southern Gangdese margin. The volcanic rocks distributed in north of Qiangtang are often regarded as part of the Hoh Xil zone, but their petrologic character, age, and spatial distribution distinguish them from Hoh Xil volcanics, requiring to be termed as &northern Tibet volcanic zone*. Typically, these include the rocks at Bamaoqiongzong, Fenghuoshan, Nianquan, Jianchuan, whose isotopic ages are mainly c. 27 每29 Ma with possible upper and

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Qinghai 每 Tibet plateau volcanic rocks occurred along adjacent regions of the Altyn strike 每 slip fault zone, including Dahingliutan, Pulu, Ashikule, Xiongyingtai, Jingyuhu, Mozitage in southern Xinjiang Province, and Hanxia, Hingliuxia, and Yumen in Gansu Province. Their radiometric ages are mostly between 0.067 and 1.4 Ma (Pleistocene) and possibly date back to the Pliocene (2.8, 4.6 Ma; Deng, 1991; Bureau of Geol. Miner. Res. of Xinjiang Autonomous Region, 1993). These eruptives are referred to as the Yutian 每Yumen volcanic zone. According to the paired model, unexposed muscovite/two-mica granites beneath the Lesser Himalaya are implied.
Fig. 11. Paired zones of the Qinghai 每 Tibet 每 Himalaya intracontinental orogenic igneous rocks (after Deng et al., 1994d, 1996a). Shoshonite series volcanic rocks belt: 1〞northern Tibet; 3〞Kekexili; 5〞Yutian 每 Yumen. Muscovite and two-mica granite belt: 2〞southern Gandise; 4〞higher Himalaya; 6〞lower Himalaya (hypothetical). Continental block: TL〞Tarim; BS〞Beishan; QC〞Qilianshan 每 Qaidam; BK〞Bayankela; QT〞Qian Tang; GD〞Gandise; HM〞Himalaya; ID〞India. Boundary: ASF〞 Altyn strike fault; NQS〞north of Qilian suture; SKS〞south of Kunlun suture; TJS〞Tuotuohe 每 Jinsajiang suture; BLS〞Bangonghu 每 Lanchangjiang suture; YZS〞Yaluzhangbujiang suture; MCT〞main central thrust fault; MBT〞main boundary thrust fault.

5.2. Orogenic episodes and horizontal growth Stages of magmatic activities characterize the orogenic episode (Deng et al., 1994c,d). Oligocene, Miocene, and Pleistocene shoshonitic series paired with muscovite/two-mica granites appear to reflect three orogenic episodes contributing to the Qinghai 每Tibet 每 Himalaya orogen. Such episodicity is consistent with the results of the British每 China Joint Expedition Team of Tibet Plateau Geology (1990) which yielded evidence on the basis of tectonic deformation for three (45 每30, 30 每5, 5 每0 Ma) episodic events. Fig. 11 shows the migration of orogen boundaries both northwards and southwards with time, reflecting lateral orogen growth along a north 每 south axis. The first episode occurred in the Oligocene, with southern and northern boundaries corresponding to the Yarlung 每 Zangbojiang and Tuotuohe 每 Jinsha Jiang sutures (northern Qiangtang block), respectively. This episode is considerd to be responsible for the formation of the Gangdese 每 Qiangtang orogen. The second episode developed mainly in the Miocene, with southern and northern boundaries corresponding to the MCT and southern Kunlun suture, respectively, being considered critical to the formation of the Qinghai 每 Tibet 每 Himalayan orogen. The third episode, occurring mainly in the Pleistocene, with southern and northern boundaries corresponding to the MBT (Main Boundary Thrust) and Altyn fault, respectively, is considered to be the main stage of formation of the present Qinghai 每 Tibet 每 Himalaya orogen. However, lateral growth of the orogen is hardly to be explained by the Indian subduction model

lower limits of 38 每39 Ma (later Eocene) and 10 Ma (early Miocene), respectively (Lai et al., 1996; Deng et al., 1994d; Deng, 1989; British-China Joint Expedition Team of Tibet Plateau Geology, 1990; Pan et al., 1990). The muscovite/two-mica granites form small stocks along the southern Gangdese; they are referred to as the &southern Gangdese zone* which have typical occurrences in areas including Xietongmen, Quewa, Xuegula, Yangbajing, Bomi and Chayu, with isotopic ages of c. 23 每 35 Ma (Deng et al., 1994d; Liu et al., 1990). The Hoh Xil volcanic zone consists of Miocene (Nl) eruptives within the Bayan Har block of the northern Qinghai 每 Tibet plateau. Recent petrologic and isotopic age studies (Deng, 1989) and detailed mapping (Sun, 1992) show that they formed c. 14 每 24 Ma, distinguishing them from the Oligocene eruptives of the northern Tibet volcanic zone (Deng et al., 1994d). The higher Himalaya and Laguiganhri muscovite/two-mica granites are mainly 10 每 23 Ma in age (Deng et al., 1994a,b) and were paired with the Hoh Xil volcanic zone. Since the Pleistocene (Q), northern

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but may be explained by the model of India 每 Tarim convergence (Deng et al., 1996a). 5.3. Asymmetry of the orogenic boundary It is notable that northern and southern margins of the orogenic belt are characterized by shoshonitic volcanism and muscovite/two-mica granite intrusive activity, respectively. Why is there such a big difference between northern and southern margins, and what does it tell us about orogenic boundary properties and deep process effects? If, as discussed above, muscovite/two-mica granites imply the northward intracontinental subduction and hotter crust coupled with colder mantle (as confirmed by recent deep seismic reflection data; Zhao and Nelson, 1993), then shoshonitic volcanism along the northern margin suggests lateral shortening or distributed thickening as a result of lithosphere convergence and colder crust coupled with hotter mantle relative to the southern margin. Recent deep seismic reflection profiles across northern Qilian to the Hexi Corridor and western Kunlun to Tarim (Gao et al., 2001; Xiao et al., 2001) suggest that Tarim 每 Beishan lithosphere has not significantly underthrust the Qilian 每 Qaidam 每 western Kunlun block and may be regarded to be formed by the distributed thickening. Thus, in contrast to the southern orogen margin, the absence of intracontinental subduction appears consistent with petrologic indications of deep mantle melting. Therefore, the differences in igneous rock association observed between northern and southern margins of the Qinghai 每 Tibet 每 Himalaya orogen suggest fundamental and asymmetric tectonic differences. 5.4. Suture zone types The higher Himalaya MCT is an intracontinental crustal suture between the overlying Tibet block and underlying Lesser Himalayan slab as it lacks evidence of ophiolite sequences. Given that collisions occur between disparate continental blocks, each one has its own margins with thinned crust and a relatively independent geological history. Such sutures may be referred to as being of &higher-Himalayan* type. Another type of continental crustal suture characteristically incorporates ophiolites, as represented by the

Yarlung 每Zangbo section, and may be referred to as &Yarlung每 Zangbo* type. Following the &hard* India 每 Asia collision, it is suggested that the Indian continent continued subducting beneath Gangdese such that the crustal suture is obscured by the overlying ophiolite suture. Accordingly, higher Himalaya and Yarlung每 Zangbo types of crustal suture reflect intracontinental block collision and continent 每 continent collision, respectively. 5.5. Crust convergence and mountain root formation Although lateral extrusion may result in horizontal crustal shortening (Tapponnier et al., 1986), it does not produce vertical thickening. On the basis of &double crust* models, the extent of horizontal shortening has been estimated at about 1400 km. Deng et al. (1995a) showed that preexisting basic crustal material was probably reequilibrated as eclogite during mountain formation prior to being incorporated by delamination into the asthenospheric mantle. Furthermore, Dewey (British 每 China Joint Expedition Team of Tibet Plateau Geology, 1990) estimated that eroded crust amounted to c. 20, 10, and 2 km in the Himalayan, Gangdese, and Kunlun regions, respectively. Thus, the amount of shortening during multiple orogenies may be absorbed by recycling both to the mantle and by erosion, being not considered in previous studies. If c. 10 km thick of crust with a width of 1400 km has vertically been absorbed, it requires about 350 km of shortening (assuming crustal thickness of c. 40 km). As indicated by Dewey (British 每 China Joint Expedition Team of Tibet Plateau Geology, 1990), crustal shortening in the regions north to the Qinghai 每 Tibet plateau must also be considered. Average crustal thickness of Tarim, Tianshan, and Altai are c. 48每 53, 50 每55, and 47 km, respectively (An et al., 1993). Assuming an average thickness of 50 km, thickening of the crust is about 10 km if the 40 km thickness of original crust is assumed. For an average width of 1200 km of the regions north to Qinghai 每 Tibet plateau, it is estimated that about 300 km shortening occurred. Thus, from the above, the total crustal shortening was about 2050 km. Given the shortening estimate of 2400 km based on palaeomagnetic studies, the 350 km deficit may be attributed to the effects of eastward crustal extrusion.

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5.6. Mantle convergence and lithosphere root formation The India 每 Tarim convergence appears to be driven by deeper mantle downwelling, whereas lateral expansion is driven by the growth of orogenic lithosphere roots. However, the presence of magmatism on the orogen margins suggesting local thermal anomalies associated with tectonically weaker zones must also be reconciled with mantle convergence effects, i.e., intracontinental subduction in the south and distributed thickening in the north (Fig. 12). Assuming overall heterogeneity of the lithosphere, two conditions are required to be satisfied during intracontinental sub-

duction: (1) while the base of the subducting continental block slopes towards the subduction zone, a trench-like feature develops between subducted and overlying blocks; and (2) the core and margin of the subducted block comprise platform-type and passive margin-type sedimentary rocks, respectively, whereas the overlying block represents a hotter active plate margin; both the higher density of the subducted block controlled by its lower temperature and sediment-derived H2O-rich fluids are needed for the intracontinental subduction to be successful. While this condition appears to be satisfied to the south of the Zangbo suture, where intracontinental subduction has occurred, this appears not to be the case for sutures between each two of the Gangdese, Qiangtang, Bayan

Fig. 12. Cartoon showing the Qinghai 每 Tibet 每 Himalaya intracontinental orogenic igneous rocks and orogenic processes (after Deng et al., 1996a). Orogenic episode: A〞Oligocene; B〞Miocene; C〞Pleistocene. Shoshonite series volcanic rock belt in north margin: a〞Northern Tibet, b〞Kekexili, c〞Yutian 每 Yumen; Muscovite and two-mica granite belt in south margin: d〞southern Gandise, e〞higher Himalaya, f〞 lower Himalaya; MCT〞main center thrust fault; MBT〞main boundary thrust fault; 1〞normal thick crust; 2〞thickened double-continental crust; 3〞shoshonite series volcanic rocks; 4〞muscovite and two-mica granite.

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Har, Qaidam, and Tarim blocks which are essentially characterized by distributed thickening. Recent study (Deng et al., 2001a,b) presented that the Tibetan lithosphere is of heterogeneity and composed of three types: Pamirs type of cool lithosphere roots, Nianqingtanggula type of thinned lithosphere, and Qiangtang type of &warm* lithosphere roots (formed by cooling of the asthenosphere), corresponding to the early-, middle- and late-tectonic phases, respectively. 5.7. Recently rapid uplift of the Qinghai 每 Tibet plateau Rapid uplift of the Qinghai 每 Tibet plateau was probably associated with the last (post-Pleistocene) orogenic episode. In contrast to the two preceding episodes, the latest orogeny was characterized by the beginning of orogenic collapse in the interior (i.e., Gangdese), while compression developed at the margins with the formation of north 每 south-trending extensional basins and lithosphere thinning from about 200 to 120 km in thickness. A possible explanation is that, during horizontal growth of the orogen, negative buoyance of the lithospheric root presumably caused its rupture and detachment from the lithosphere. The resulting inflow of asthenosphere into the delaminated root zone would cause a reduction in mantle density, and thus, the resulting buoyance effect led to rapid uplift of the lithosphere. However, because the orogen margins are considered to be affected by both compressional stress and mountain root buoyance, the rapid uplift resulted from a combination of these mechanisms.

6. Mantle upwelling and lithosphere stretching During the Cenozoic, both East Asian continental rifting and western Pacific marginal sea formation were clearly dominating features in the global context. Here, continental rifting in eastern China is considered in relation to the mantle plume-like upwelling. 6.1. Continental rift magmatism The distribution and petrological characters of volcanism in eastern China indicate two main epi-

sodes of basalt magmatic activity: Paleogene and Neogene 每Quaternary. The former was largely associated with two large rift-related basins, i.e., the lower Liao River, north China, Bohai, Subei, south Yellow Sea, and Jianghan basin zone in the west, and the East China Sea and South China Sea regions in the east. Neogene 每 Quaternary activity occured mainly in smaller rifted basins and is concentrated in the following three zones from west to east: Daxinganling 每 Taihangshan, Changbaishan 每 Tan Lu fault, and southeast coastal zones (Deng, 1988). Eastern Chinese continental rift-related magmatic rocks consist mainly of tholeiite and alkali basalt with bimodal (basic to acid) differentiates, indicating variably LREE- and LILE-enriched sources (Deng, 1988; cf. Condie, 1982). The Cenozoic basalts are richer in HFSE with average contents of TiO2, 2.25 wt.%, Nb, 58 ppm, and Zr, 285 ppm, being typical for rift-related basalts. Primary or near-primary magmas are mostly alkali basalts, although Paleogene basalts in the Beijing area are tholeiitic. This pattern resembles the Red Sea每 East Africa rift associations and is strongly contrary to Yanshanian (J 每 K) calc 每 alkaline, HKCA, and shoshonitic character in eastern China. The bimodal character is dominated by mafic compositions with minor felsic eruptives (trachyte 每 comendite) appearing only in Tianchi of Changbaishan (Deng, 1988). Magmas from rifts with larger extensional velocity produced alkali basalt (Ne-norm = 0 每 5%), minor basanite (Ne-norm = 5每 9%), and locally, hy-normative tholeiite. Rift flank eruptions formed mostly basanite (Ne-norm z 9 每 10%), and, locally, nephelinite, with a lateral zonation resembling that of the East African rift. The spatial variation of Neogene 每Quaternary magmas clearly relates to rifting rather than oceanic subduction with compositional porality (Table 8). The Paleogene basalts show higher contents of K2O and K2O + Na2O near the coast than inland (Table 8), possibly reflecting shallower magmatic sources beneath north China than beneath Bohai 每 Subei. 6.2. Crust 每mantle thermal structure Temperature distribution is a significant parameter in discussing crust 每mantle dynamics. Eastern Chinese Neogene 每Quaternary basalts contain abundant

J.F. Deng et al. / Earth-Science Reviews 65 (2004) 223每275 Table 8 Average K2O and total alkali contents in basalts (%; after Deng, 1988) Neogene 每 Quaternary Daxinganling 每 Taihangshan zone West flank K2O K2O + Na2O Axis East flank 2.18 5.78 Changbaishan 每 Tanlu fault zone West flank 2.18 5.78 Axis 2.00 5.60 East flank 2.37 6.03 Southeast coast zone West flank 2.13 5.55 Axis 1.75 4.96 East flank 2.16 5.75 0.80 1.22 3.47 3.87 Inland p coast Paleogene North China

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Bohai 每 Subei

1.59 1.19 5.10 5.04 Inland p coast

ultramafic xenoliths, including peridotites and minor pyroxenites, from which a pyroxene geothermal gradient was first interpolated by Deng et al. (1980). This suggested that the Cenozoic upper mantle thermal state beneath eastern China resembled that of oceanic regions indicating upper mantle diapirism. The &anomalous* thermal state of eastern China has probably been an important factor during Cenozoic tectonic reactivation (Deng et al., 1980, 1987). By 1990, a total of 405 pyroxene thermobarmetric estimates allowed precise definition of the regional geotherm and corroborated the rift-related perturbation of the upper mantle thermal gradient (Deng et al., 1990b, 1991, Deng et al., 1996a). Here, we discuss the inferred crust 每mantle thermal structure beneath eastern China in the context of rapid Neogene每 Quaternary lithosphere thinning. Normally, the crust 每mantle thermal state shows a curved distribution produced by the upward change from convective to conductive heat transfer in the Earth*s thermal boundary layer. However, Morgan (1983, 1984) have shown that heat transfer in the lithosphere became dominantly convective during active tectonism and magmatism due to perturbation of the lithospheric thermal structure. Thus, during rapid lithosphere thinning, the quasiequilibrium thermal curve will be deflected producing, in Eastern China, the type of thermal structure depicted in Fig. 13. In this model, the solid line beneath the Moho represents an average geothermal gradient of 3.3 jC/km, similar to the gradient of a dry peridotite solidus (about 4 jC/km). Taking end-members of the adiabatic geotherm (0.5 jC/km) represented by the convective heat transfer model and average conductive geotherm (14 jC/km, derived from Pollack*s global model), the contributions of both convective and conductive heat transfer to the

eastern Chinese upper mantle geotherm (3.3 jC/km) have been computed to be 79% and 21%, respectively (Deng et al., 1990b, 1991, Deng et al., 1996a). The dashed line below the Moho (Deng et al., 1996a) in Fig. 13 becomes smooth near the depths of about 75 km contoured from the maximum density of pyroxene thermobarometric estimates (after Deng et al., 1990b, 1991). This depth may correspond to the average depth of the lithosphere 每 asthenosphere boundary (L/A in Fig. 13). Above this boundary, the lithospheric mantle geotherm was calculated to be

Fig. 13. Thermal structure of the crust/upper mantle in east China. The solid line is made by Deng et al. (1990b, 1991); the dashed line is after Deng et al. (1996a). The curve above Moho is by Wu (1994), The curve near Moho is by Deng et al. (1990b, 1991). L/A is lithosphere/asthenosphere boundary.

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about 5.0 jC/km based on the estimation of relative contribution of convective and conductive heat transfer being about 67% and 33%, respectively. Below the boundary, the asthenospheric geotherm was calculated to be about 2.0 jC/km, closer to the adiabatic gradient, with convective and conductive heat transfer components being estimated at about 89% and 11%, respectively (Deng et al., 1996a). Although the thermal structure of continental crust is genarally regarded to reflect conductive heat transfer, the geotherm derived from early Precambrian metamorphic rocks in northern China becomes smooth at depths of about 10 km. Below this depth, the middle to lower crustal geotherm (7.3 jC/km) is steeper than predicted by the conductive heat transfer model (20 jC/km), with convective and conductive contributions being estimated to be about 65% and 35%, respectively. This probably reflects tectonically induced thermal perturbations in the middle and lower crust. A widespread hydrous fluid-bearing, low-velocity, high-conductivity layer at present-day midcrustal levels in northern China may be an effective environment for convective heat transfer. Accordingly, Wu (1994) proposed a new model, suggesting that present-day geothermal distribution beneath eastern Hebei Province is dominated by convective heat transfer in the middle crust and by conductive heat transfer in the upper and lower crust. This model predicts that

temperatures at the top of the middle crust may reach 500 jC, in contrast to only 250 每350 jC in the depth of 10 每 15 km predicted by the conductive model. Because the 500 jC isotherm is regarded as a boundary between brittle and ductile continental crustal rocks (Wu et al., 1994), the new model is able to explain the concentration of earthquake focal depths at 10 每 15 km beneath northern China, being consistent with depth of the Curie isotherm. On the basis of theoretical considerations, Morgan and Baker (1983) suggested &rapid* and &slow* endmember models for thermal thinning of the lithosphere (Fig. 14). Fig. 14a shows a model for slow rates of thinning, in which the solid line 1 represents the thermal structure before thinning. During lithospheric thinning, the upper surface of the asthenosphere ascends from L/A (1) to L/A (2). If the thinning rate is very slow, heat from the asthenosphere will be conducted to the lithosphere base, and the temperature at point P will increase to that at q. Thus, temperature changes in the overlying lithosphere represent a quasiequilibrium state. After thinning, the thermal structure is represented by the dashed line 2, on which the segment above q is the geotherm reflecting conductive heat transfer while the segment below q represents the near-adiabtic convective asthenospheric geotherm. The combined effects of lithosphere thinning and asthenospheric heating will

Fig. 14. Hypothetical thermal structure model of slow (a) and rapid (b) lithospheric thermal thinning (after Morgan, 1983). L/A〞lithosphere 每 asthenosphere boundary: L/A (1) before lithosphere thinning, L/A (2) after lithosphere thinning. Corresponding geotherm: solid line 1 before thinning and dashed line 2 after thinning. Dot of P and q is described in detail in the text.

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produce regional uplift. If the thinning ceases, regional uplift will also cease, and the uplifted region will be characterized by a heat flow anomaly. Fig. 14b shows a model for rapid rates of thinning (cf. Fig. 14a), in which the temperature at P (base of the lithospheric after thinning) does not match a rapid temperature increase, and the lithosphere thermal structure is essentially the same as that before thinning. This geotherm is shown as dashed line 2 and represents the thermal structure of the lithosphere and asthenosphere in the brief period after thinning. When thinning ceases, the lithosphere becomes heated, and the temperature at its base ( P) increases to the temperature at q (Fig. 14b). Eventually, a quasiequilibrium thermal structure is obtained, indicated by the dashed line in Fig. 14a. This lithospheric heating process is referred to as thermal relaxation. A consequence of rapid lithospheric thinning is that two stages of uplift occur. In the first of these, uplift is rapid and accompanied by a local thermal anomaly resulting from magmatic activity with no regional heat flow anomaly. In the second stage, uplift occurs after the end of thinning and is related to thermal relaxation. When thermal relaxation ends, regional uplift ceases, and a regional heat flow anomaly results. Thus, guided by Morgan*s models, the crust 每 mantle thermal structure of eastern China may be clarified. The structure shown in Fig. 13 is similar to that in Fig. 14b and clearly suggests that thermal thinning in eastern China was rapid. We infer that the observed disequilibrium (i.e., perturbed) thermal structure is the product of a tectonothermal event which accompanied continental rifting. Deng et al. (1987) proposed that mantle diapirism was the main cause of the Moho thermal inflection, whereas a secondary thermal kink at the upper/middle crustal boundary resulted from diapirism of mid- or lower crustal material and thermal convection induced by midcrustal fluid circulation. Accordingly, we conclude that the crust 每 mantle thermal structure of eastern China (Fig. 13) represents a record of rapid thermal thinning of the lithosphere. Remaining problems include questions concerning the stage of thermal evolution represented by eastern Chinese lithosphere and the extent to which future thermal evolution of the region may be predicted. We can reasonably assume that the correspondence of the upper boundary of the asthenosphere inferred from

geophysical and basalt thermobarometric data represents the end of a thermal thinning event. Two major episodes of Cenozoic thermal thinning are recorded in eastern China by widespread voluminous basalt eruptions in the Paleogene and Neogene 每 Quaternary, respectively; the latter includes historic activity in regions such as Changbaishan and Wudalianchi. Thermal thinning induced by the Neogene 每 Quaternary episode may have progressed to thermal relaxation. This is supported by the observation that the eastern part of China is still being uplifted today (Gui et al., 1989), and that heat flow anomalies are confined to sedimentary basins and volcanic regions, with heat flow elsewhere being more or less normal. It is predicted, however, that lithosphere heating and regional uplift will continue in the future, with expansion of the area showing anomalous heat flow. The lithosphere will only begin to cool towards the end of the thermal relaxation stage before eventually reaching a stable thermal state. 6.3. Asthenosphere petrology The nature and depth of asthenosphere can be deduced and constrained by the compositions and genetic conditions of primary (e.g., mantle xenolithbearing) basaltic magmas. On the basis of experimentally determined mineral-melt phase equilibrium, we calculated thermodynamic P 每T equations for magmas using available data on activities of the chemical components involved. The cross point of several P 每 T curves is the P 每T equilibrium condition between magma and peridotite in the upper mantle, representing the condition of magma segregation from its mantle source (Carmichael, 1977; Deng, 1987). Assuming H2O-free conditions, the segregation depth of mantle-derived primary basaltic magmas was estimated to be in the range of 50每 100 km, mostly between c. 50 and 80 km, i.e., 50 每100 km (mainly 60 每 80 km) in northern and northeastern China, 43 每56 km in Hannuoba, 55 每 72 km in northern Jiangsu, 50 km in the coastal region of southeastern China, 65 km along the Xinchang 每 Mingu zone, and 100 km along Juxian每 Mingxi zone, respectively (Deng et al., 1990a). For H2O-present conditions, estimated depths are somewhat deeper because the presence of H2O decreases the temperature and SiO2 activity of basaltic magma. For an average content of

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1 wt.% water, the segregation depth was estimated to be c. 5 每 6 km deeper than that of dry magmas, implying that water-bearing magmas in eastern China segregated from depths in ranging between c. 55 and 85 km (Deng et al., 1990a). Geophysical data indicate that upper boundary of the upper mantle low-velocity layer ranges between c. 60 and 80 km beneath eastern China. For example, Pwave velocities (Vp) indicate 60每 68 km (Song et al., 1986), and S-wave velocities (Vs) indicate 60 每 80 km (Song et al., 1992), whereas deep penetration seismic refraction data indicate 83 km beneath Xingtai (Teng et al., 1982). Magnetotelluric surveys show that the lowvelocity layer is 60 每80 km beneath the north China plain (Liu, 1985). The heat flow values of 84 mW/m2 in the Paleogene and 64 mW/m2 at the present time indicate 55 and 75 km of the asthenospheric top, respectively (Deng et al., 1990a). The Paleogene boudinage extensional model suggests that the top of the asthenosphere has a depth of 60 km (Deng, 1988). The concurrence of these estimates implies that low-velocity layer temperatures exceed those of the upper mantle solidus, implying the presence of small interstitial melt fractions (Table 9; Deng et al., 1985). If the upper asthenosphere boundary in the Cenozoic was the same as at present, basalt segregation depths may be assumed to be similar. Although it is difficult to infer past asthenospheric depths, the igneous petrologic probe method, based on estimated magma segregation depths and melt fractions, may provide important thermal constraints on the formation and evolution of eastern China rift zones (Table 9). Logachev (1983) proposed that melt fractions in the upper asthenosphere may be 0.1 beneath Kenya rift and 0.05 beneath the Baikal rift, corresponding to Vp values of 7.5 and < 7.7 km/s, respectively. The low-velocity
Table 9 Nature of the upper asthenosphere inferred from basaltic magma generation (after Deng, 1988; Deng et al., 1996a) Time Pliocene 每 quaternary Miocene Paleogene Depth to asthenosphere 50 每 85 km, (average 70 km) 65 每 100 km, (average 85 km) 50 (Beijing) 每 55 km (Bohai, Northern Jiangsu), (average 53 km) Amount of interstitial magma 3 每 10%, (mainly 6 每 8%) 3 每 10%, (mainly 5 每 7%) 15 (Beijing) 每 18% (Bohai, Northern Jiangsu)

layer Vp values beneath the north China plain and northern China are 7.2 每 7.4 and 7.6 km/s, respectively (Teng et al., 1982; Song et al., 1981), implying, according to the petrologic probe method, melt fractions in the range of 0.07 每0.15 (Table 9; cf. Logachev, 1983). The geophysical results thus support our model and affirm the importance of combining the petrologic probe approach with data from geophysical studies. 6.4. Lithosphere dynamics Lithospheric extension, thinning, and uplift are interrelated aspects of continental rifting, representing shallow-level and surface responses to processes in the mantle. Although the development of normal faults and grabens in eastern China during the Cenozoic indicates crustal extension, it is not clear whether these features completely penetrate the lithosphere. We propose, however, that the large volume basaltic eruptions represent evidence favoring complete penetration, assuming that these originated in the upper asthenosphere and propagated to the surface via extensional fractures. 6.4.1. Lithospheric extension Lithospheric extension is generally unstable. Neugebauer (1983) pointed out that there are two types of lithospheric deformation which is caused by upwelling of mantle material, i.e., boudinage (necking) tectonics and doming tectonics, of which the first is generally the more common. It appears that in eastern China, boudinage tectonics developed during Paleogene and was succeeded by doming tectonics in the Neogene 每Quaternary (Deng et al., 1992b). 6.4.1.1. Boudinage tectonics. Large-scale boudinage tectonics are typified by the Basin and Range Province of western North America. Froidevaux and Ricard (1985) proposed a multiple-layer boudinage model in which: 4H < k < 8H (k is wavelength, i.e., the width of basin or range; H is the thickness of boudinage layer). They pointed out that the surface topography in western North America (k = 50 km) exhibited the deformation in upper crust, and the variation of Bouguer anomalies (k = 200 km) exhibited boudinage of the lithosphere. Paleogene &basin and range* tectonics occurred in eastern China as

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Fig. 15. Schematic profile of Paleogene basin 每 range tectonics.

shown schematically in Fig. 15. The k values for north China basin, Northern Jiangsu basin, Hetao basin and Zhegan basin were approximately 300, 258, 50 and 30 每40 km, respectively. The k values between the uplifts of Hetao and north China basins and between the uplifts of north China and northern Jiangsu basins are about 400 and 350 km, respectively (after Fig. 3 in Wang et al., 1983). The northern Chinese basin and range tectonics exhibits a k value of 60 km (estimated from data in Xu et al., 1985). Taking k = 6H, the lithospheric thickness is 60 km (k = 350 km) beneath the north China and northern Jiangsu basins, and upper crustal thickness is about 10 km beneath the north China, Hetao, and Zhegan basins. The estimated lithospheric thickness coincides with the upper asthenosphere boundary referred to above (Section 6.3). Estimated upper crust thickness is consistent with the upper boundary of low-velocity crust (12 每16 km) and the lower limit of listric faults (5每 17 km; Deng et al., 1992b). These results indicate that both large-scale and secondary-scale variations of topography express boudinage tectonics of the lithosphere and upper crust, respectively. The absence of Paleogene basalt in Hetao and Zhegan basins implies that extension was limited to shallow crust. 6.4.1.2. Doming tectonics. A second phase of lithospheric extension occurred in eastern China during the Neogene 每 Quaternary. Large graben-faulting basins, which developed in the Paleogene, were filled by sediment and then subsided. However, small-scale faulted basins became widespread within uplift zones flanking the large Paleogene basins and were sites of intense basalt volcanism. Although the surface topography still exhibits basin and range character, and the small-scale basins may reflect boudinage tectonics in

the upper crust, the large-scale topography does not mirror the shape of the upper asthenosphere surface. However, both surfaces appear to exhibit the same topography which is indicative of doming tectonics (Fig. 16). Given that magmas here originated at depths of c. 60每 85 km, this suggests that the uppermost asthenosphere surface beneath the uplift zone had changed from concave in the Paleogene (Fig. 15) to a convex shape in the Neogene每 Quaternary (Fig. 16). 6.4.1.3. Extension types and lithospheric viscosity Neugebauer (1983) studied the relationship between extension induced by mantle diapirism and the excess viscosity of crust (Fig. 17). The excess crustal viscosity parameter (g2/g1) refers to the excess of crustal viscosity over that of the subcrustal lithosphere, where g2 and g1 represent the viscosity of crust and lithospheric mantle, respectively. The g2/g1 value of the Chinese continent is 1每 10 in the Paleogene when large boundinage tectonics developed, changing to 103 每 104 in the Neogene每 Quaternary where the lithosphere exhibited doming tectonics. 6.4.2. Extension type and the rate of thermal thinning Theoretical model calculations (Morgan and Baker, 1983) show that the viscosity of lithospheric mantle

Fig. 16. Schematic profile of Neogene 每 Quarternary lithospheric doming-type extension deformation.

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Fig. 17. Relationship between the types of extension and the parameter of excess of crustal viscosity referred to the subcrustal lithosphere (after Neugebauer, 1983). L/A〞lithosphere 每 asthenosphere boundary; g1 and g2 are subcrustal lithosphere and crustal viscosity, respectively.

(g1) needs to decrease to 1020 每 1022 Pa?s to allow asthenospheric diapirism over time spans of tens of millions of years. Although data for the viscosity of lithospheric mantle in the Paleogene are as yet unavailable, the available viscosity data for the Neogene 每 Quaternary lithospheric mantle (Deng et al., 1988) permit at least qualitative discussion of relationships between lithospheric extension type and viscosity (Fig. 17). Studies of peridotite xenoliths in Neogene basalts from Hannuoba, Hebei Province yield g1 value for lithospheric mantle of 1020 每 1022 Pa?s, for depths < 75 km, and 1019 每1020Pa?s for the asthenospheric mantle at depths >75 km (Deng et al., 1988), the former being consistent with the theoretical value. Cenozoic magmatism in eastern China was almost exclusively basaltic with little or no eruption of crustderived rhyolitic magma (an exception being the Quaternary Changbaishan volcano), suggesting that the crust was cold. This inference supports the assumption that crustal viscosity (g2) has not changed since the Paleogene. Accordingly, and assuming the g2/g1 values in Fig. 17, we infer that g2/g1 in the Paleogene (characterized by boudinage tectonics) should be smaller than that in Neogene 每Quaternary (characterized by doming tectonics). In other words, the Paleogene mantle g1 value would be greater than that in the Neogene 每 Quaternary. However, as the Paleogene asthenosphere was shallower than that during the Neogene每 Quaternary (Table 9), Paleogene g1 should be less than that in the Neogene 每Quaternary (assuming that the rate of thermal thinning remained unchanged). This contradiction may be reconciled if the Paleogene rate of thermal thinning were faster than

that in the Neogene每 Quaternary, given that more rapid thinning would minimize or preclude lithosphere heating. This is consistent with its lower temperature and higher viscosity. In this case, the Paleogene g2/g1 value would be smaller and correspond, as expected, with boudinage tectonics. From the above discussion, we conclude that, if the lithospheric thermal structure in the Neogene 每 Quaternary reflects rapid thermal thinning, Paleogene must represent a record of extremely rapid thermal thinning. On this basis, and in consideration of Morgan*s model, we propose a new theoretical model for the three categories of thermal thinning rate (Table 10). Therefore, the type of thermal thinning can be interpreted from the type of extension, uplift event, and distribution of heat flow anomalies. This model provides a useful framework for inversion of the relationship between deep-seated processes and their shallow responses.

Table 10 Types of lithospheric thermal thinning (after Deng et al., 1996a) Types Extensional Uplifting event (magma) type Large boudinage tectonics Large basins; (mantlederived basalts) Two phase uplift, (1) main uplift accompanied with thermal thinning; (2) uplifting accompanied with thermal relaxation ditto Heat flow Example anomaly Local Paleogene anomalies in eastern China

Extremely rapid thermal thinning

Regional anomalies

Rapid Doming, thermal small scale thinning upper crust boudinage; rift basins (mantlederived basalts) Slow tectonic thermal features as thinning above; (significant mixing of crust-mantle magma)

ditto

Neogene 每 Quaternary in eastern China

One-phase uplift〞 terminated by thermal thinning

Regional Yanshanian heat flow (Jura 每 anomaly Cretaceous) in eastern China

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The case for slow-rate thermal thinning depicted in Table 10 is still tentative and requires further study. Widespread mantle-derived, crust 每 mantle-mixed, and crustal source Yanshanian (J 每K) magmatism occurring in eastern China indicates that the whole lithosphere was heated and showed regional heat flow anomaly characteristics. The several uplifting events accompanying magmatism and the rapid erosion and peneplanation events (unconformities) correspond to the uplifting during thermal thinning and the end of uplifting after the cessation of magmatism, respectively. These petrologic records may be taken as an indication of slow-rate thermal thinning of the lithosphere. 6.4.3. Lithosphere thinning and uplift The process of lithosphere thinning includes mechanical thinning by extension, thermal thinning, and thinning caused directly by asthenospheric diapirism. The latter two are not readily distinguished as both are related to ascending of asthenosphere. Here, we refer to both, together, as thermal thinning. 6.4.3.1. Mechanical thinning. Mechanical thinning of the lithosphere is produced by extensional deformation. Based on the observed deformation in faulted basins, an average extensional rate of f 15% was inferred for the north China plain during the Paleogene and f 5% during the Neogene 每Quaternary (Xu et al., 1985). Compared to present-day thickness of lithosphere (60 km) and crust (35 km) based on geophysical studies, inverted thickness for the beginning of the Neogene and the Paleogene (prior to continental rifting) are 63 and 37, and 74 and 44 km, respectively. The value of 44 km is consistent with the 46每 47 km crustal thickness observed for stable Erdos. Therefore, Cenozoic extensional deformation produced c. 14 km lithosphere thinning and c. 9 km crustal thinning. It should be pointed out that deformation of fault basins is limited to the brittle upper crust. Because extensional thinning in ductile middle and lower crust is almost certainly greater than in the upper crust, the above estimate for mechanical crustal thinning should be taken as a minimum value. This is supported by deep seismic studies of north China basins (Sun et al., 1988), indicating attenuated middle and lower crustal layers compared with surrounding areas. Lithospheric thickness is also con-

trolled by changes in depth of the upper asthenosphere surface, i.e., thermal thinning. Thermal thinning does not affect crustal thickness because it is confined within the lithospheric mantle. Therefore, calculated crustal thickness probably represents a lower limit prior to the advent of rifting. If a prerift thickness of c. 44 km is representative, eastern Chinese crust had recovered isostatically to its normal thickness by the late Mesozoic or early Cenozoic. REE distribution patterns of Quaternary Changbaishan trachytes exhibit pronounced negative Eu anomalies (Tang and Tian, 1989) in distinct contrast to the Yanshanian trachytes without negative Eu anomalies (Deng et al., 1996a), indicating normal thickness or thinning of the Cenozoic crust. 6.4.3.2. Thermal thinning and uplift. Morgan (1983) and Neugebauer (1983) pointed out that crustal uplift (1每 2 km in general) is a widespread feature of rifted continental terrains representing the isostatic response to thermal thinning. Mechanical thinning contributes only to subsidence of the rift base and is not responsible for regional uplift. The Paleogene eastern Chinese peneplain surface is expressed by topographic height of 2200 每3700 m a.s.l. (Ma, 1985). Miocene basalts are currently located in upland regions of volcanic rift basins, whereas Quaternary basalts occupy lower elevations and valleys. The numerous valley terranes exhibit signs of continuous crustal uplift during the Cenozoic. Geodetic measurements in northern China over 20 years (Gui et al., 1989) show that the north China plain continues to be uplifted at an average rate of 3 每 5 mm/yr, reaching 20 mm/yr locally. Assuming that widespread, continuing uplift of eastern China is an isostatic response to lithospheric thermal thinning, the extent of uplift and densities of lithosphere and asthenosphere may be estimated. According to Fig. 18, assuming the boundary of lithosphere and asthenosphere (L/A) before thinning is an isostatic surface, it can be shown that: H ? qL ? ?H ? qA ? ? ?h ? qL ? h=H ? ?qL ? qA ?=qL where qL and qA are the densities of lithosphere and asthenosphere, respectively; h is the amount of crustal

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Fig. 18. Domal uplift induced by the lithospheric thermal thinning (after Deng et al., 1988).

by thermal thinning (H) and mechanical extension (Deng, 1988; Table 12). From another approach, Zhang et al. (1983), Menzies et al. (1993) and Griffin et al. (1998) presented that the lithosphere had been thinned about >120 or 80 每140 km, respectively, by a comparison between the Archean lithospheric root of 180 每220 or 200 km and the Cenozoic lithosphere of 75 每80 or 60 每120 km in thickness. 6.5. Asthenosphere dynamics

uplift, and H is the extent of lithospheric thinning (Deng, 1988). Considering the early Paleogene peneplain to be currently at 2200每 3700 m a.s.l. (Ma, 1985) and those for the Neogene 每 Quaternary at 1200 每400 m a.s.l., the magnitude of uplift in the Paleogene (h) can be taken as 2 km. The base of Miocene basalts in the Hannuoba 每 Datong area occurs at 1350 m a.s.l., and the base of Pleistocene basalts occurs at 1050 m a.s.l. The late Paleogene peneplain has therefore been uplifted by 400 每 500 m (Deng et al., 1985). The assumption of 700 m post-Neogene uplift (h) and a qL value of 3.3 g/cm3 (Ringwood, 1975) are taken. Experimental q values for olivine-bearing tholeiitic and alkali basaltic magmas at 1.5 GPa are 2.82 and 2.77 g/cm3, respectively (Kushiro, 1982). For respective melt fractions of 0.15 and 0.07 for Paleogene tholeiites and Neogene 每 Quaternary alkali basalts (Tables 9 and 11), estimates of qA in the Paleogene and for the Neogene 每Quaternary are estimated to be 3.25 and 3.29 g/cm3, respectively (Table 11). Using the above values and formula, we can calculate the thickness of lithosphere &converted* into asthenosphere by thermal thinning (Deng, 1988; Table 11). Taking the average thickness of modern lithosphere as 60 km, we can also deduce the lithosphere thickness before rifting (HL) after adding lithosphere removed
Table 11 Estimated partial melt fraction ( F ), thickness, and density of lithosphere and lithosphere transferred to asthenosphere (after Deng, 1988) qmagma (g/cm3) Paleogene Neogene 每 Quaternary 2.82 2.77 F 0.15 0.07 qL (g/cm3) 3.33 3.33 qA (g/cm3) 3.25 3.29 h/km 2 0.7 H/km 83 58

6.5.1. &Plumes* and &subplumes* Oceanic island and intraplate continental basaltic volcanism have long been regarded as the surface expression of mantle plumes or &hotspots*, whereby hotspots weaken the lithosphere allowing for decompression melting and excess volcanism (Wyllie, 1984; Anderson, 1981; Morgan and Baker, 1983). In northern China, there were large volumes of the Cenozoic rift-related basaltic eruptions associated with lithosphere extension, rapid thermal thinning, and uplift, being consistent with the notion of a sublithospheric mantle plume (Deng et al., 1992a,b,c,d). White and McKenzie (1989) considered the uppermost parts of such plumes to have diameters of c. 1000每 2000 km and show temperatures c. 100每 200 jC greater than normal. Deng et al. (1998b) paid attention to the proximity of four western Pacific marginal seas, the South China Sea, Sea of Japan, Sea of Okhotsk, and Bering Sea, to the major regions of Cenozoic riftrelated basalt eruptions in East Asia which was interpreted as possible surface manifestations of discrete sublithospheric &plumes*. We suggest that sufficient data exist in support of convective mantle upwelling beneath northern China
Table 12 Estimated lithosphere thickness (HL) prior to rifting (after Deng, 1988) Thickness (km) after Thickness (km) after HL/km mechanical thinning thermal thinning Before the 74 Paleogene rifting 63 Before the Neogene 每 Quaternary rifting 83 157

58

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to permit discussion of its formation and structure. The surface expression of upwelling mantle appears to be ellipsoidal with long- and short-cell diameters of about 1800 and 1350 km, respectively, comprising &subplumes* measuring 450 每 525 km (500 km on average) in diameter (Fig. 19). &Subplumes* are here defined as second-order features that develop from the primary (&first-order*) plume. These are shown schematically in Fig. 19, each &subplume* circled by the

distribution of related Cenozoic basalts (Deng et al., 1992a,b,c,d). During the Paleogene, the plume would have comprised four subplumes (I, II, III, and IV) corresponding to the Sea of Japan region, the Bohai 每 North China region, Shuangliao 每lower Liaohe River, and Mudanjiang, respectively. In the Neogene每 Quaternary, the primary plume also consisted of four subplumes (V, VI, VII, and VIII) corresponding to the Changbaishan, Zhangjiakou 每 Abag Qi, Great

Fig. 19. Mantle plume and sub-plume beneath north part of the China continent and the adjacent regions (after Deng et al., 1992c, 1998b). 1Exposures of Cenozoic basalt; 2-Surface expression of sub-plume and releated volcanic areas. Paleogene: I-Japan sea, d = 450 km, ocean crust; II-Bohai-North China plain, d = 450 km, sodium volcanic rocks; III-Shuangliao-Xialiaohe plain, d = 450 km, sodium volcanic rocks; IVMudanjiang, d = 450 km, potassic volcanic rocks. Neogene-Quaternary: V-Changbaishan, d = 525 km, sodium and potassic volcanic rocks; VIHannuoba, d = 525 km, sodium volcanic rocks. VII-Daxinganling, d = 450 km, sodium volcanic rocks; VIII-Wudalianchi, d = 450 km, potassium-rich volcanic rocks.

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Hingganling, and Wudalianchi volcanic regions, respectively. 6.5.2. Origins and forms of mantle upwelling To discuss the formation of mantle plumes in this context, it is necessary to define relationships between temperature distribution and mantle melting. Although the pyroxene geotherm derived from basaltborne mantle xenoliths (Deng and Zhao, 1991) may reflect upper mantle thermal structure during basalt volcanism, it is only valid to depths of about 160 km. It may, however, be extrapolated to c. 400 km along an adiabatic geotherm of 0.5 jC/km (Ringwood, 1975; Fig. 20). This precept may be tested through other approaches. For example, the modified Carmichael activity method suggests that, under anhydrous conditions, basalts are generated at 2 每 3 GPa and 1300每 1500 jC, whereas hydrous basalts are generated at 2 每3 GPa and 1200 每 1400 jC (assuming melt contains c. 1% H2O; Deng et al., 1992a,b,c,d).

Basalt eruption temperatures, calculated from meltmineral geothermometers to be c. 1130 每1200 jC, allow interpolation of melt segregation temperatures of about 1200 每 1250 jC at 2 每 3 GPa assuming adiabatic magma ascent (Deng et al., 1992a,b,c,d). Melt segregation temperatures of about 1270每 1300 jC at 2 每3 GPa may be interpolated from one atmosphere experimental data for the xenolith-bearing Datong basalt whose liquidus temperature is 1237 jC (Qin et al., 1994). Validity of the pyroxene geotherm (Fig. 20) is corroborated by three independent approaches. The H2O content of the mantle source is estimated at about 0.1% as interpreted from 1% H2O contents of basaltic magmas and assumed ambient melt fractions of about 0.1. In Fig. 20, pyrolite solidi for varying amounts of H2O and CO2 (Wyllie, 1988) are used as a basis for modeling mantle melting. Dry and H2O-excess peridotite solidi are shown for comparison (Wyllie, 1988). The pyroxene geotherm clearly intersects solidi of both 0.1% H2O-bearing pyrolite and the peridotite 每 C 每 H 每 O system at depths of about 75 km, indicating the upper limit of partial melting in mantle peridotite, being consistent with asthenospere depths of 60每 80 km (Deng and Zhao, 1990a,b). Intersection of the pyroxene geotherm with the peridotite solidus occurs at two pressure conditions, and the higher pressure is equivalent to about 400 km, representing the lower limit of partial melting (Fig. 20). Between c. 400 m and 75 km in depth, the pyroxene geotherm lies above the solidus implying increased melt fractions up to about 100 km depth. These decreases thereafter with melting terminated at about 75 km at the base of rigid lithosphere. It is suggested therefore that plume-related melting commenced at depths of about 400 km in the presence of volatiles released by deeper mantle degassing. Density perturbations resulting from partial melting would reinforce mantle plume ascent, whose rate increases with progressive increases in melt fraction. 6.5.3. Petrologic structure of mantle plumes Petrologic studies of Cenozoic volcanism in eastern China (Deng et al., 1992a,b,c,d) indicate that there are two primary magma types, xenolith-bearing alkali basalts, which are commonly erupted to the surface, and picritic tholeiites, which intermittently reach the surface due to their higher density. The latter are,

Fig. 20. A pyroxene geotherm beneath northeast China and the peridotite and pyrolite and pyrolite solidus with 0.1% H2O (after Deng et al., 1992b,c); 1〞pyroxene geotherm; 2〞shield geotherm; 3〞pyrolite solidus with 0.1% H2O; 4〞solidus of peridotite 每 C 每 H 每 O system; 5〞dry peridotite solidus; 6〞H2O-oversaturated peridotite solidus. Data from: 1, Deng and Zhao (1991); 3, Ringwood (1989); 2, 4, 5, and 6, Wyllie (1988).

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however, represented by their less dense fractionation products among surface eruptions. It might be inferred from this observation that mantle plume upwelling comprises two parts, a central part characterized by picritic interstitial melt and marginal parts characterized by alkali basaltic interstitial melt (Fig. 21). 6.5.4. Thermal structure of mantle plumes Primary magma fields and pyroxene geotherms are schematically shown in relation to the litho-

sphere 每 asthenosphere boundary (L/A) and projected isotherm distribution (Fig. 21). This depicts the structure and shape of a hypothetical mantle plume comprising: (1) a central part with temperatures up to c. 1500 jC, and interstitial picritic tholeiite melt flanked by (2) lower temperature regions with smaller alkali basaltic melt fractions. Such a plume is interpreted to have ascended to relatively higher levels in the Paleogene than during the Neogene 每 Quaternary.

Fig. 21. Schematic section of the shape and structure of the mantle plume beneath northern China in (a) Paleogene and (b) Neogene 每 Quaternary (after Deng et al., 1992c,b). The number represents temperature of the isotherm. The ruled region denotes the plume region containing picritic thoeleiitic interstitial melt; L/A is lithosphere/asthenosphere boundary.

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6.5.5. Shape of mantle plumes We suggest that during the Paleogene, MORB-like oceanic crust formed during opening of the Sea of Japan above the eastern part of the plume, whereas, above the western part of the plume, rift-related basalts were erupted in northern China. It is also conjectured that, at earlier stages, the central part of the plume underlay the Sea of Japan, with a diameter (at deeper levels) of only 400 每 500 km. On rising to the base of the lithosphere, the plume developed a mushroom-shaped head in response to the resistance of rigid lithosphere. Thus, as shown in Figs. 19 and 21, the surface expression of the plume center probably corresponds to the Sea of Japan (I in Fig. 19), and the mushroom-shaped head may underlie the Bohai 每 north China plain, Shuangliao 每lower Liaohe River, and Mudanjiang volcanic regions (II, III, and IV in Fig. 19). Kinematic relationships (Figs. 19 and 21) suggest that the plume head diameter (long and short diameters being 1800 and 1350 km, respectively) is larger than that of the plume center (400 每 500 km). Interestingly, we note that the plume forms an asymmetrical &umbrella* extending northwestwards, in contrast to the hypothetical model of White and McKenzie (1989), a probable result of its confinement by Pacific plate subduction. Southeastward migration of the Sea of Japan in the Neogene 每 Quaternary led to volcanism in the Changbaishan region as the plume center impinged on northern Chinese lithosphere (Figs. 19 and 21). Compared to other northern Chinese Neogene 每 Quaternary basaltic centers, Changbaishan activity is more intense, widespread, and longlasting, and is also uniquely characterized by late Quaternary rhyolite volcanism at Tianchi. These features are consistent with mantle upwelling beneath the Changbaishan region. 6.6. Asthenosphere 每 lithosphere interactions 6.6.1. Migrating volcanism as a record of lithosphere motion The &hotspot reference frame* is often used to measure the velocity and trajectory of lithospheric plates. Fig. 19 shows that, in comparison with Paleogene activity (Fig. 19; I, II, III and IV), Neogene 每 Quaternary volcanism in northern and eastern China (Fig. 19; V, VI, VII and VIII) clearly shifted north-

westwards, being consistent with the notion that Chinese lithosphere moved southeastwards relative to the proposed plume center. The positions of corresponding Paleogene and Neogene 每 Quaternary &subplumes* indicate that volcanism shifted 525 每 675 km to the northwest (Fig. 19), and that Chinese lithosphere drifted c. 600 km to the southeast. It is, however, more difficult to calculate plate motion velocities because of their time dependence. In the latest Oligocene to early Miocene, large rift-basins affected by Eocene to Oligocene basalt volcanism were uplifted and eroded producing major unconformities (Bureau of Geol. Miner. Res. of Hebei Province, 1989). In uplifted regions beyond these basins, Miocene basalts unconformably overlie Mesozoic or preMesozoic strata as in Wangqing (Jilin), Hannuoba (Hebei), Linqu and Penglai (Shandong), Penglai (Hainan Island), Mingxi (Fujian), and Xinchang and Shengxian (Zhejiang) (Deng, 1988). In Mongolia, the Miocene Daligange basalts directly overlie the Oligocene 每Miocene peneplain surface. Prior to the inception of Baikal rifting, the Late Cretaceous 每Paleogene weathered crust persisted up to the latest Oligocene 每 early Miocene (Deng, 1988). Data for the earliest Miocene basalt eruptions are restricted to a few locations, including detailed paleontological data on the Jinlongkou (Cixian) and Xuehuashan (Jingxing) formations, which indicate a mid-Miocene lower limit (Bureau of Geol. Miner. Res. of Hebei Province, 1989). The geological record shows that the period from Oligocene to early Miocene may be regarded as a gap between the two stages of volcanism and rifting. Isotopic dating for Cenozoic basalts in northern China yield only 10 early Miocene ages (24 每17 Ma) among a reasonably definitive range of 116 K 每 Ar dates (Liu, 1992). An average lithosphere motion rate of 8.57 cm/a between c. 2 and 17 Ma was estimated, being comparable to 9.66 cm/a for Hawaii and to 6.6, 8.6, and 9.5 cm/a for the Indian, Pacific, and Nazca plate motions, respectively (Liu, 1985). Southeastward lithosphere drift and fan-shaped opening of the Japan Sea in the early Miocene can explain several critical features including: (1) plumeinduced Paleogene volcanism in rift-related basins, such as the Bohai, north China, and north Jiangsu basins, and the Sea of Japan, terminated by their removal from underlying plume influence; (2) Neo-

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gene 每 Quaternary rift-related volcanism initiated in the mid-Miocene in response to lithosphere motion over the plume; and (3) early Miocene fan-shaped opening of the Japan Sea (21 每 14 Ma) evidenced by counterclockwise rotation of NE Japan and clockwise rotation of SW Japan (Tatsumi et al., 1990). 6.6.2. Lateral volcanic zonation as a record of lithosphere rifting The symmetrical zonation of rift-related magma in composition between rift flanks and basins (Condie, 1982; Deng et al., 1985) is comparable to that observed at mid-ocean ridges (Deng et al., 1984; Figs. 22 and 23). Pleistocene basaltic eruptions (so-called &valley basalts*) are confined to low-lying basins, while the Pliocene basalts were erupted on the elevated basin perimeters. Miocene basalts are located in high mountains or plateaus surrounding the basins. The latter two categories are referred to as &high-level* basalts. However, the Pliocene basalts at higher elevation are in direct contact with pre-Cenozoic strata which lack intercalated Miocene or Pleistocene basalts. Pleistocene basalts directly overlie Quaternary

Fig. 23. Distribution of basalts in Jingyu 每 Jingbohu vocanic basin (after Deng, 1988). 1〞N1 basalt; 2〞N2 basalt; 3〞Pleistocene basalt.

Fig. 22. Distribution of basalts in Hannuoba 每 Datong volcanic basin. 1〞N1 basalt; 2〞N2 basalt; 3〞Q1 basalt; 4〞Q2 basalt (after Deng, 1988).

sediments which in turn overlie pre-Cenozoic, including Archean, formations. The distribution of basalt types in graben basins thus reflects a &stair-shaped* pattern, indicating that the rifted basins resulted from lateral spreading, whereby progressively older volcanic edifices are thrust aside. It also indicates a general pattern of crustal uplift (Fig. 24). Considering the lateral shift of volcanism (Figs. 22 and 23), the extension rates in the Hannuoba每 Datong and Jiyu 每 Jinbohu basins are estimated to be 0.22 and 0.23 cm/yr, respectively, based on the relationship: v (cm/yr) = S/Dt (Deng, 1988). These values are consistent with extension rates of 0.20 cm/yr derived from the inversion of basalt composition (Deng, 1985) and comparable to the extension rates of up to 0.30 cm/yr estimated for the east Mrican rift (Deng et al., 1993). This broadly linear volcanic belt was broken into two symmetric belts resulting from horizontal spreading between which the original extensional axis lay prior to its disruption. According to Figs. 22 and 23, the axis of Quaternary volcanoes lay to the northwestern side of the Neogene volcanic axis in both the Hannuoba 每Datong and Jinyu 每 Jinbohu basins, implying southeastward lithosphere drift relative to the linear subplume magma source. The calculated aver-

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Fig. 24. Model for the formation of continental rift 每 volcanic basin (after Deng, 1988).

age drift rates are 0.04 and 0.05 cm/yr, respectively, for these two basins (Deng et al., 1993). If the Pleistocene volcanoes in Jinyu and Jinbohu are considered, the axial line trends NE 45j, indicating a 15j rotation from the c. NE 30j Pliocene axis. An estimated clockwise rotation rate of 0.76 ? 10? 6j/yr can explain why Neogene basalts were not erupted to the northwest of Quaternary activity in the Jinbohu area. Rift spreading was often asymmetric (Fig. 25) because of the superimposed effects of lithosphere drift on horizontal extension. 6.6.3. Volcanic &&riveting** of the lithosphere? The rate of lithosphere drift during active volcanism is apparently very low, only 0.04 每 0.05 cm/a (Deng, 1988), such that drift rates of c. 8.57 cm/a during a volcanic hiatus are two orders of magnitude

greater. We ascribe this effect to a &riveting* of the lithosphere, whereby volcanic activity effectively impedes its lateral motion (Deng et al., 1992a,b,c,d). Sublithospheric mantle plumes or subplumes may serve as multiple &rivets*, whereby their enhanced coupling with and penetration of the overlying lithosphere results in significant slowing of lithosphere drift. Accordingly, absence of the &rivet effect* during volcanic quiescence allows relatively free motion of the lithosphere.

7. Dynamic evolution of the Chinese lithosphere In this section, the evidence of crust and mantle structure, magmatic activity, and orogenic accretion are synthesized into a geodynamic model for the evolution of Chinese continental lithosphere. 7.1. The assembly of China 7.1.1. Collision orogeny: the main mechanism As for other continents, eastern Eurasia comprises a number of stable blocks separated by orogenic belts. Those making up the Chinese lithosphere are notably smaller, however, highlighting China as a &field laboratory* amenable for studying supercontinent assembly. Geological and paleomagnetic studies show that stable Chinese continental blocks show discrete kinematic and geological histories prior to their assembly. Moreover, an overwhelmingly significant part of the record is preserved in orogens resulting from the progressive accretion of continental plate fragments.

Fig. 25. Asymmetrical spreading model (after Deng, 1988).

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Orogenies resulting from continental plate collisions represent the terminal stage of repeated cycles of ocean basin opening and closure (the Wilson Cycle). In general, orogenic &sutures* incorporate much of the (albeit fragmented) geological record of a Wilson Cycle, especially that relating to preexisting ocean basins. Ophiolite melanges have long been identified with the latter and are common features of orogens. For example, the Yarlung Zangbo ophiolite marks a suture produced by the &hard* collision of Indian and Eurasia at c. 45 Ma. The accompanying mountainbuilding episode is referred to by some as a &postcollision intracontinental* orogeny, and by others as a &continent 每continent* collision orogeny. We propose the existence of two suture types, a &Yarlung Zangbo* type and a &higher Himalayan* type (see above). The first of these is characterized by the presence of ophiolites, representing paleo-ocean remnants entrapped by a continent 每 continent collision. The second suture type lacks ophiolite and is interpreted to represent colliding continental terrains. Accordingly, ophiolitic sutures are not necessarily exclusive indicators of collision orogenies. It appears, in fact, that orogens produced by intracontinental terrain collisions (i.e., &intracontinent collision orogens*) are more widespread than &continent 每continent*-type orogens. A collision orogeny results in the formation of a continental root, roughly double the thickness of normal crust. Continental roots can result from either of two processes, superposition of discrete continental crust sections or vertical thickening by horizontal compression and shortening. We refer to the former process as an &intracontinental subduction* collision, characterized by muscovite/two-mica granites, and the latter as a &horizontal shortening* collision, characterized by shoshonitic igneous activity according to which they can be readily distinguished in the geological record. 7.1.2. The Indosinian stage As a basis for understanding collision orogenies, we synthesized geochemical data for Chinese muscovite/ two-mica granites and also for A-type granitoids which characteristically terminate an orogeny (see Fig. 26). Triassic muscovite/two-mica granites in the southern part of Beishan (Gansu Province) were displaced westwards by the left-lateral slip of Altyn Tagh fault in the Cenozoic. Restored to their preslip locations,

these granites would have been located in the eastern part of China (east of 100jE), suggesting that eastern China was largely assembled in the Indosinian. Radiometric ages for the granites range from 254 (Zuo, 1992, unpublished report) to 244 Ma (Li, 1993, personal communication). A bifurcating muscovite/ two-mica granite belt also occurs in eastern Inner Mongolia and northernmost China emplaced between the Early Permian and Late Jurassic. The northern branch shows U 每Pb and K 每 Ar ages of 253 每 222 Ma (Bureau of Geol. Miner. Res. of Inner Mongolia Autonomous Region, 1991; Luo et al., 1995), whereas the southern branch shows ages of 264.7 每 199 Ma (Bureau of Geol. Miner. Res. of Shanxi Province, 1989; Bureau of Geol. Miner. Res. of Hebei Province, 1989; Bureau of Geol. Miner. Res. of Inner Mongolia Autonomous Region, 1991; Shi, 1994, personal communication). Late Triassic muscovite/two-mica granites in Tengchong 每 Lincang, (Yunnan Province) extend southwards into Burma, Thailand, and Malaysia and show ages of 237每 211 Ma (Bureau of Geol. Miner. Res. of Yunnan Province, 1990). Muscovite/ two-mica granites from Triassic to Cretaceous with isotopic ages of c. 245 每122 and c. 225每 113 Ma, respectively, along northern and southern margins of the Yangtze craton (see Fig. 26) were described in detail by Deng et al. (1995b). Indosinian collisional reactivation of the stable north China platform is evidenced by a major unconformity separating upper Triassic from middle Proterozoic to Early 每 Middle Triassic formations (Wang, 1996; Zhao, 1990; Cui et al., 2000). An Early to Middle Triassic basin developing in the north Chinese interior was initially connected to the ocean via south Qilian and Ganzi (Ma, 1992). However, this cover fold was truncated by the overthrust &Inner Mongolia basement axis* along east 每 west Shangyi 每 Chifeng 每 Longhua 每Jianping 每 Beipiao in northernmost China. The overthrust &Inner Mongolia axis* has Triassic muscovite/two-mica granites, and, to the south, the foreland fold 每thrust belt of northern margin of the north China platform is located (Fig. 27). The southdirected thrust belt consists of ductile-deformed mylonite hornblende Ar 每 Ar ages of c. 211 Ma (Wang, 1996). Late Triassic molasse and the evidence of thrust fronts clearly indicate northward intracontinental subduction of the north China block (Fig. 27), analogous to that interpreted for the subduction of

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Fig. 26. Scheme of distribution of muscovite/two-mica granites and A-type granites on the China continent since Permian (after Deng et al., 1996a). 1〞Muscovite/two-mica granites; 2〞A-type granite; 3〞eclogite. The English letter of each type represents its formation era.

Greater India at the Himalayan MCT. Assuming that the Indosinian in north China comprise foreland or Jura-type folding, the &Inner Mongolia basement axis* belongs to the Hingan 每 Mongolia orogen, exhumated

Fig. 27. Model of the intracontinental orogeny of the north China continent subducting northwards during the Indosinian (after Deng et al., 1996a). 1〞Foreland folded belt; 2〞Inner Mongolia crystalline basement; 3〞cover rock; 4〞muscovite and two-mica granite; 5〞the direction of north China continent subduction.

in Indosinian, rather than to that of the north China platform. The unconformity between Late Jurassic volcanic rocks and crystalline basement suggests the latter was exposed towards the end of the Middle Jurassic. The absence of ophiolite from the Indosinian belt supports the conclusion that it represents a &higher Himalayan*-type orogen. A further question is whether the Indosinian system is a unique orogenic feature, a continuing episode of the Hingan 每Mongolia Hercynian orogeny or the initial phase of the eastern Chinese Yanshanian orogeny. In general, orogenies are terminated by extensional collapse and the appearance of A-type granitoid magmas (see Section 4). From Fig. 26, we know that Permian A-type granitoids were widespread in the Junggar 每 Hingan 每 Zhangguangcai Hercyjan orogen. Studies of ophiolite belts between Hegenshan in the east to Souloushan每 Karamery and west Junggar in

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the west show Early to Middle Devonian ages. To the south, a belt extending from Soulunshan to Bayingou (north Tianshan) may have formed in the early Carboniferous (Xiao et al., 1992) as is also the case for the Hongshishan 每 Baiheshan ophiolite belt in Beishan (Gansu Province; Deng et al., 1996b). Limited data for muscovite/two-mica granites from the Altyn mountains and north Tianshan, c. 331.9 Ma (Zou et al., 1988), south Mongol Gobi, Late Devonian (Zuo et al., 1990), and Inner Mongolia, 375 Ma (Tang and Zhang, 1991) indicate that a Late Devonian to Carboniferous intracontinental subduction orogeny followed oceanic closure, then followed by postorogenic collapse in the Permian. The subsequent Triassic orogeny was a distinct &intracontinental* rather than &continent 每continent*-type orogeny. Given that the end of the Indosinian orogeny is marked by widespread Late Triassic A-type granites (Fig. 26), the Yanshanian orogeny was clearly an independent phenomenon which was also confirmed by A-type granitoid aged 229 每 202 Ma (Heilongjiang; Hong et al., 1991), 218.8 and 235.9 Ma (Fanshan and Yangyuan), 234 Ma (Liyuanhekanzi), 240每 190 Ma (Saima), 224 Ma (Antuqinglizi; Zou et al., 1988), 185.8 and 203 Ma (Ulongshan, Fengning, and Guangtoushan每 Pngquan; Wang et al., 1994), and 248.8每 219.2 Ma (Luojingou in Tianzhen, Shanxi Province; Bureau of Geol. Miner. Res. of Shanxi Province, 1989). In contrast to northern China, Late Triassic A-type granites appear to be absent from the south China and Songpan 每 Ganzi orogens, respectively, marking eastern and western margins of the Yangtze continent (Fig. 26). This suggests that Indosinian and Yanshanian orogenies in south China represent successive &intracontinental* rather than &continent 每 continent* collisions. The ultrahigh pressure coesite eclogites in Dabieshan and Sulu have the Sm 每Nd isochron ages of 244每 221 Ma (Li et al., 1989), but coeval muscovite/two-mica granites are absent. This indicates that the Indosinian collision orogeny between the northern and southern continents in the eastern China was achieved by horizontal shortening by means of lithosphere subduction combined with strike 每slip shearing (Deng et al., 1995b). From this discussion, it is concluded that the Indosinian assembly of eastern China was largely accomplished by intracontinental as opposed to continent 每 continent collisional oroge-

nies. Assuming, therefore, that continental subduction was the dominant process, horizontal shortening and strike 每 slip faulting may be taken to accommodate the bulk of convergence between southern and northern continental blocks. 7.1.3. The Yanshanian stage The occurrence of Jurassic and Cretaceous muscovite/two-mica granite belts to the west of 100jE (Fig. 26) reflects assembly of Central China in the Yanshanian. From north to south, the Jurassic granitoid belts are as follows: (1) northern Qaidam belt which extends from Eboliang to Rongkawanyin, including the Tatalenhe pluton, with K 每 Ar ages of 200 每163 Ma and intruding Middle Triassic rocks (Bureau of Geol. Miner. Res. of Qinghai Province, 1991); (2) Kunlun belt extending from Yutian to Idatan and Nachitai, with radiometric ages of c. 174 Ma for Yutian (Xu, 1994, personal communication), 194 Ma for Xidatan and 198 Ma for Nachitai plutons; and (3) Bayanhar belt extending from Zhajia to Zhawulong (Bureau of Geol. Miner. Res. of Qinghai Province, 1991). Cretaceous muscovite/two-mica granite belts (Fig. 26) include: (1) Nianqing Tangula belt in the west extending from Bange to Naqu, radiometric ages being 108.5 Ma (Duola pluton), 99 Ma (Sangxing pluton; Li et al., 1982); (2) an eastern Tibet belt extending along the Leiwuqi to Xiaya and Zuogong with ages of 134 每79 Ma (Chengdu Institute of Geol. & Miner Res. 1989, unpubl. report); and (3) west Yunnan belt, located in Biluoxueshan to Luxi, further extending southwards into Burma, Thailand, and Malaysia with ages of 112 每68 Ma (Bureau of Geol. Miner. Res. of Yunnan Province, 1990). Some of these, e.g., the Kunlun, Bayanhar, east Tibet, and west Yunnan belts, are associated with coeval ophiolite belts, whereas others, e.g., the northern Qaidam and Nianqing Tangula belts, are not. However, all muscovite/two-mica granite belts are related to late Indosinian (northern) or late Yanshanian (southern) Tethyan closure episodes (Deng et al., 1994c). In general, these reflect continent 每continent orogenies and, in some cases, were accompanied by intracontinental collisions where intracontinental subduction dominated. Orogenic mechanism in eastern China clearly differs from those in western China. In general, the former were associated with the inner continental margin orogeny which was related to subduction of

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the Izanaqi ocean (Deng et al., 1996c, 1999, 2000a; Davis et al., 2001). We have already argued that the Yanshanian Songpan 每 Ganzi and south China orogenies succeeded the Indosinian intracontinental orogeny, and that a continental margin-type orogeny characterized southeastern coast of China. The Yanshanian orogeny in the southern part of eastern China thus reflects the combined effects of oceanic subduction and intracontinental block collision. However, orogenic mechanisms are more complex in the northern part of eastern China. A large-scale continental margin orogeny in the Jurassic overlapped with intracontinental block collision orogenies, the latter recorded by the east Hebei muscovite/two-mica granite belt dated at 166 每171 Ma (Wang et al., 1994; Ye et al., 1991). This extends northeastwards and connects with the Yanji pluton belt (208 每 170 Ma; Zhou et al., 1992) and so-called gneiss每 granite series (206 每183 Ma; Lin et al., 1992) in east Liaoning and Shandong. It is noted from Fig. 26 that Jurassic muscovite/twomica granite belts occur in the Sulu region, an extension of the Dabie ultrahigh pressure metamorphic terrain, implying that Sulu intracontinental subduction is consistent with indentor tectonics (Yin and Nie, 1993). Late Early Cretaceous A-type granites in eastern China and Korea mark the extensional collapse of orogens and termination of Yanshanian intracontinental orogeny. Radiometric ages for typical alkaline granites are: 134 Ma (Laoshan in Qingdao; Lin et al., 1992), 91.3 Ma (Quiqi, Fujian), 115 每 70 Ma (Foguoshi, Korea; Hong et al., 1987), 118 Ma (Houshihushan), 131 Ma (Wulingshan), and 100 Ma (Xiangshan, all in east Hebei; Wang et al., 1994). When Early Cretaceous postorogenic collapse occurred over most of eastern China, muscovite/two-mica granite belts developed in the Wandashan region of northeasternmost China at 137.5每 106.1 Ma (cf. Deng et al., 1996a). These granites record continent 每continent or arc 每 continent collisions following the closure of the Mesozoic Mongolia 每 Okhotsk 每 Wandashan ocean. Accordingly, we conclude that the main part of China was bounded by oceanic subduction zones to the east and southwest. In general, Mesozoic orogenic activity expanded from north to south, as Chinese continental growth occurred from north to south around Siberia. However, the northeast-trending Early Cretaceous Atype granites cut east 每 west-aligned Late Triassic A-

type granites, highlighting distinct Indosinian and Yanshanian orogenic stress fields. 7.1.4. Crustal and lithospheric roots Petrologic evidence from igneous rocks supporting the presence of crustal and lithospheric roots (discussed in Sections 4 and 5) in the Mesozoic is complemented by studies of ultrahigh pressure metamorphic lithologies in Dabie 每Sulu and Inner Mongolia. As proposed earlier (Section 3), thickened Qilianshan lower crust probably reached granulite facies conditions. Crustal root compositions are granitic &eclogite*, while the upper mantle lid, a component of the lithospheric root, is probably &basic* eclogite. While none of these interpolated lithologies is exposed at the surface, high-pressure granitic and basic granulites in the Inner Mongolia axis are known to have equilibrated at 1.2 每1.5 and 1.4 每1.5 GPa, respectively (Qian and Wang, 1994), which are consistent with the interpolated depth range of Qilianshan lower crust (see Section 3), suggesting that the petrologic collision 每 response model is a reasonable prospect. Moreover, studies of the Dabie ultrahigh pressure (UHP) terrain suggest coesite-bearing eclogite dikes and sills equilibrated under garnet amphibolite to amphibolite P 每 T conditions. Zircons enclosed by Dabie eclogites (in turn, enclosed by marble, coesite pseudomorphs by garnet, and accompanying jadeitic quartzite) yield multipleformation ages (c. 2000, 800 每 1000, 400 每500, and 220 Ma; Yan, personal communication, 1985) which include effects of the Indosinian orogeny. Magnesitebearing garnet dunites also yield equilibration depths of c. 150每 210 km (Yang et al., 1993), implying roots extending to at least c. 210 km in depth. UHP lithologies formed at pressures between 1.6 and 3.0 GPa (Suo et al., 1993) correspond to those of the present-day Qilianshan crustal root as inferred from geophysical data. Given that the Dabie 每 Sulu UHP lithologies are not a product of oceanic subduction and are attributed to horizontal shortening during continental lithosphere subduction (Deng et al., 2000b), coesite-bearing eclogites are clearly predicted in the Qilianshan upper mantle lid. Thus, the records of exhumed UHP metamorphic terrains strongly support the proposed model for the evolution of crustal and lithospheric roots in response to intracontinental orogenies.

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7.2. Lithosphere dynamics The prime question raised in regard to Chinese lithosphere dynamics is the extent to which a genetic relationship exists between circum-Pacific tectonics in the east and the Tibet 每 Himalaya orogen in the west. While some workers emphasized the distinction in tectonic orientation and style between the two domains, others believe that extensional tectonics in the east are directly related to the indentation of Asia by India, and (in some cases) the southward indentation of Siberia. For the most part, however, attention has been focused on one or other region rather than possible interrelationships between the two. The notion of extrusion tectonics associated with the indentation of India has been widely used to explain tectonic processes in eastern China (Tapponnier et al., 1986). However, England (1985) showed that, according to the numerical simulations of a thin viscous sheet model, the distance influenced by a compressional boundary is equivalent only to the length of the plate boundary itself, implying that transfer of deformation to Siberia by the India 每 Asia collision is an unlikely possibility. An alternative possibility, suggested by Willett and Beaumont (1994), is that Asia has been subducted beneath Tibet. One-directional indentation could explain the formation of South China Sea but not that of the Japan Sea along with horizontal growth of Tibet and Himalayan orogen. Deep subcontinental crust and mantle processes have gained increasing attention in recent years (Deng et al., 1996a; Flower et al., 1998). The slip-line field theory of Tapponnier et al. (1986) cannot explain the formation of lithosphere root and mountain root in west China and the formation of lithospheric thinning and mantle plume*s upwelling in east China. Is there some coupling relation between deep processes within the crust and mantle and shallow-level lithosphere tectonics and magmatic activity? This question is now discussed from a three-dimensional (3-D) perspective. 7.2.1. Double indentation 每extension tectonics: a 2-D planar model 7.2.1.1. Southward indentation of Siberia. While effects of the northward indentation of India are widely recognized, those relating to possible south-

ward indentation of Siberia are less so. However, a series of southward convex orogenic arcs were well developed to the south of the Siberian block supporting the notion of indentation by the block. From north to south, these comprise the eastern Sayan 每Baikal arc, Irkusk arc, Mongolia arc, and Qilianshan 每 Luliang arc (Liu, 1980). Liu (1980) referred to the trace of the tip of these arcs as the &north每 south belt of East Asia* (Helan 每Kangdian line), pointing out its correspondence to the Angora 每 Tunguse fracture system in middle Siberia. To the east of this line, the main tectonic orientation is northeastwards, while to the west it is northwestwards. Northeast- and northwesttrending faults exhibit sinistral and dextral slips, respectively, a kinematic pattern that appears to confirm the southward indentation of Siberia. Accordingly, we (Deng et al., 1996a) emphasize that convex orogenic arcs abutting both the northern margin of India and southward margin of Siberia are strong indications that the tectonic deformation of eastern China is a product of both effects. 7.2.1.2. Eastward escape of China. The distribution and slip directions of faults in central Eurasia and their related seismicity are consistent with the slip 每 line field predicted from the indentation of &rigid wedges* such as Siberia, India, and Arabia. Here, we briefly discuss the tectonic framework induced by Siberian indentation and eastward escape of East Asia as defined by the 2-D planar indentation model of Tapponnier et al. (1986) (Fig. 28). The left-lateral Mongol 每Okhotsk fault in Fig. 29 can be taken as the I2K2K1 fault in the mirror figure of Fig. 28 (Liu, 1980). The northern margin fault of north China showed right-lateral slip in the Tertiary (Wan, 1993) and corresponds to the F1 fault in the mirror figure (Fig. 28). Together, the two faults accommodated eastward extrusion of the Mongol 每 NE China block (corresponding to the B1 block in Fig. 28a). This process may be linked to opening of the Japan Sea (corresponding to the wedge between B1 and B2 blocks in Fig. 28b). In the Neogene 每 Quaternary, the slip sense of the northern margin fault reversed to left lateral (Wan, 1993) and, together with the northern margin fault of the Qinling 每Dabieshan orogen, accommodated the eastward extrusion of northern China, terminating Japan Sea opening.

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Fig. 28. A summarized result of the indentation experiments (after Tapponnier et al., 1986). Penetration of the indenter in (a) was about 2 cm, and 6.5 cm in (b). See text for discussion.

However, the slip sense of the Qinling northern margin fault is still unclear. Of two likely possibilities, one is that left-lateral slip, like that of the Qinling southern margin fault, would allow eastward motion with displacement of north China relative to the south China and Yangtze blocks. A second possibility is that the Qinling north margin fault is right lateral, such that

the Qinling orogen is a transform zone allowing for westward motion. The combination of westward Qinling motion and eastward motion of Qilianshan acommodated by left-lateral slip on the Altyn Tagh fault would have produced an east 每 west compressional field and, in turn, north 每 south striking Riyueshan in east Qinghai and the inland Qinghai Lake (cf. Deng,

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Fig. 29. Bidirectional indentation and polyphase extrusion tectonics in Asia (after Deng et al., 1996a). Both the base map and the content of the lower part of the map are after Tapponnier et al. (1986). Boundary: (1) horizontal pushing like the bulldozer or distributed thickening; (2) intracontinental subduction zone; (3) oceanic subduction zone; (4) strike 每 slip fault and slip direction; (5) marginal sea basin and extensional direction; (6) indentation direction of India and Siberia; (7) extrusional direction of major block; (8) numbers on or near all arrows refer to extrusion phase: 1〞Paleogene, 2〞Neogene 每 Quaternary. Fault: M.O.F.〞Mongolia 每 Okhotsk fault; F.N.M.N.C.〞fault of north margin of north China; A.F.T.〞Altyn Tagh fault; F.N.M.Q.F〞fault of north margin of Qinling; R.R.F.〞Red River fault. Margin Sea: O.S.〞Okhotsk Sea; J.S.〞Japan Sea; S.C.S.〞South China Sea; A.S.〞Andaman Sea.

1996a). Meanwhile, the combination of westward movement of Beishan (Gansu Province) and eastward movement of north China produced an extensional stress field and induced formation of the Badanjilin Desert. Although many problems remain, the close resemblance of inferred planar deformation for Siberian indentation with that of India and its temporal correspondence with Japan Sea and South China Sea basin opening support the model for simultaneous double indentation (Deng et al., 1996a). Comparisons be-

tween the indentation experiments and the inferred geological effects are shown in Table 13. 7.2.2. A three-dimensional model 7.2.2.1. Boundary geometry and its sense. Several aspects of boundary geometry are critical to modeling regional tectonic processes. For example, the eastward extrusion of eastern Eurasia is only possible given the lack of a constraining eastern boundary. The 3-D geometry of the rigid indentor is also likely to be

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Table 13 Main events of the East Asia extrusion eastwards induced from India and Siberia bidirectional indentation (after Deng et al., 1996a) Main events in the southern part of East Asia related to the India indentation Paleogene (E)
. The Red

Main events in the northern part of East Asia related to the Siberia indentation
. The right-

In comparion to the model of the indentation experiments (Fig. 28)

.

.

Neogene(N) 每 . Quaternary (Q)

.

.

.

E 每 Q (?)

.

Source

.

. F1 lateral strike 每 slip fault of the north margin of the North China. . The . B1 The Sundaland Mongolia 每 extrusion northeast China eastwards. block extrusion eastwards . The Japan sea . The gap The South China Sea opening between B1 opening and B2 . The right. F2 The Altyn Tagh leftlateral strike 每 lateral strike 每 slip fault of the slip fault. north margin of the Qinling 每 Dabie (?). . The left-lateral . F1 The Red river rightstrike 每 slip fault lateral strike 每 of the north slip fault margin of the North China. . The North . B2 The Yangtze block extrusion China block southeastwards. extrusion eastwards. The cessation . The cessation of the seaof the sea-floor floor spreading spreading of of the South the Japan Sea. China Sea. The formation . The formation . Close of the of the Okhotsk to K1, K2 Andaman sea Sea and Baikal and Burma rift close to the lowlands close east margin of to the east Siberia margin of indentation. India indenter . Deng et al., . Tapponnier Tapponnier et al., 1986 1996a et al., 1986

River leftlateral strike 每 slip fault.

significant, although this aspect was not considered in 2-D experiments conducted by Tapponnier et al. (1986). If the front of the rigid indentor is vertical or steeply inclined with respect to the plastic body, two distinct arc deformation belts would be expected to develop. The experimental results of Tapponnier et al. (1986, figs. 3, 4, 12) indicate that where the eastern boundary is free, arc deformation in the front of the indentor is convex towards the direction of indentation. However, with regard to the indentation of India, the Himalayan range is convex southwards in sharp contrast to the northward indentation direction (Fig. 29) and the experimental indications. This discrepancy can be attributed to the 3-D geometry of the indentor (Fig. 30). Assuming the Indian front to be a plane inclined towards Asia, Asia is effectively being thrust onto the Indian plate producing the equivalent to a convex volcanic arc above an oceanic subduction zone (Fig. 30b). In contrast, we can see from Fig. 29 that arc belts to the south of Siberia are all convex to the south, similar to the direction of indentation, suggesting that the Siberian front is a vertical plane (cf. Fig. 30a) rather than inclined as is the case for India. The 3-D geometry of arc deformation in front of an inclined plane indentor (Fig. 30b) is thus consistent with continental subduction and, in turn, with the presence of muscovite/two-mica granites which (as

Fig. 30. Cartoon showing the arc structure with (a) forward convex and (b) backward convex (after Deng et al., 1996a); 1 and 2 are indenter and more plastic body of multilayered plasticine, respectively. The left and right sides represent horizontal plan and vertical cross-section, respectively.

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concluded earlier) mark intracontinental subduction boundaries. Therefore, we take arc deformation belts opposed to the indentation direction and muscovite/ two-mica granites belts as the respective deformation and petrologic records of intracontinental subduction, respectively. In contrast, the 3-D geometry in Fig. 30a reflects horizontal shortening that corresponds with shoshonitic igneous activity which marks the boundary of shortening, the respective deformation and petrologic records of horizontal shortening boundary, respectively. Considering again the indentation boundary of India, opposed southward Himalayan convexity and northward Indian indentation is matched by the presence of muscovite/two-mica granites as expected at an intracontinental subduction boundary. However, the convexity of eastern Tibetan and Pamir arcs is concordant with the Indian indentation direction, implying that they conform to horizontal shortening boundaries. This inference is supported by the absence of muscovite/two-mica granite belt in the syntaxes. Likewise, the concordance of west Kunlun southward convexity, at the northern margin of Tibet, with the direction of Tarim indentation implies a horizontal-shortening boundary, supported in turn by the presence of shoshonitic volcanism. Thus, the combined evidence of arc deformation sense and igneous activity uniquely characterizes an indentation boundary. We therefore conclude with a degree of confidence that the Indian front represents an intracontinental subduction boundary, whereas the Siberian front is a shortening boundary. The most highly compressed deformation belt coincides with Kunlun 每 Altyn Tagh 每 Qilianshan, marking the northern margin of the Tibetan plateau. The Pacific side is an oceanic subduction boundary characterized by downgoing motion and forming a relative free planar boundary. Otherwise, southward convex arc deformation in the northern part of East Asia, including west Kunlun, is not consistent with subduction of Asia beneath Tibet (cf. Willet et al., 1994). 7.2.2.2. Comparisons between east and west. Comparison of eastern and western China is constrained by limited data especially for the region between Siberia and Tarim which is not discussed further. Principal results are given in Table 14, in which geological

events since the collision of Tibet were taken from the discussion in Section 5. The timing of late Oligocene 每 early Miocene Tibetan uplift and erosion remains controversial, although many workers concur that 5000 m peak elevations represent a peneplane surface. The 1990 joint British每 Chinese expedition concluded that the surface was younger than late Eocene, and that erosion had persisted over a long period of time. Huang and Chen (1980) proposed that the Kunlun 每Hoh Xil region was covered by Pliocene 每 Pleistocene volcanic rocks, whereas later studies suggested that the volcanic rocks were Miocene (24 每 14 Ma; cf. Deng et al., 1996a) in age, leading to the conclusion that it dated from the late Oligocene to early Miocene. Its intersection with the 10 Ma Laguigangri granite indicates that peneplanation was completed by the middle to late Miocene. In Section 5, we proposed that the high Himalaya 每 Laguigangri muscovite/two-mica granite belt and Hoh Xil shoshonitic volcanics represent a paired igneous belt formed in the same orogenic episode. If this interpretation is correct, the surface beneath the Hoh Xil volcanic rocks is not the same as that cutting the Laguigangri granite, and the latter should be younger than the former. Opening of the Japan Sea occurred in two stages (Tatsumi et al., 1990; Karp and Lelikev, 1991; Shimazu et al., 1990), the Paleogene episode, in which the geometry of opening was parallel, and the early Miocene (21 每14 Ma) episode showing a fanshaped opening. The earlier stage corresponded to the formation of rifted basins in eastern China, while the second stage was coeval with the cessation of eastern Chinese basalt volcanism and eastward lithosphere drift. Later on, in the middle Miocene (17 每10 Ma), collision between blocks within the Japanese arc was accompanied by the intrusion of muscovite/two-mica granites (Fig. 26). The opening of South China Sea basin also occurred in two stages, i.e., Oligocene (32 每27 Ma) and early Miocene (23 每17 Ma) episodes (Chen, 1990). The first stage coincided with Paleogene rifting in eastern China while the second was coeval with the inception and development of the Luzon Arc. Eventual collision of the latter with Eurasia in eastern Taiwan and the north Palawan block with the Luzon arc terminated the opening of South China Sea in the late Miocene. Table 14 shows that conjugate compression and extension, orogeny and basin formation,

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J.F. Deng et al. / Earth-Science Reviews 65 (2004) 223每275 Table 14 (continued ) Qinghai 每 Tibet 每 Himalaya orogenic belt East China and adjacent areas

Table 14 Roughly comparison of the geological events on the east and west China and adjacent areas (after Deng et al., 1996a) Qinghai 每 Tibet 每 Himalaya orogenic belt
. First episode of the Paleogene (Oligocene) formation of orogenic lithospheric root . Intracontinental orogeny

East China and adjacent areas
. First episode of the

. Intracontinental

orogenic igneous activities in both the margins . Uplifting, Latest Oligocene 每 denudation and earliest planation of the Miocene orogenic belt

upwelling of the mantle plume . Formation of both the Japan Sea and the South China Sea, and large rift basins, as well as eastward extrusion of the Sundaland and Mongolia 每 northeast China block . Widespread extensive basaltic magma eruptions
. Rapid eastward

. Third episode . Since Pleistocene, Neogene 每 Quaternary (mainly in Pleistocene) weakening and finally of the formation of cessation of the mantle orogenic lithosphere plume upwelling and root at both the margins the basaltic magma erption due to less feeding of the asthenospheric materials from the west . Intracontinental orogeny only at both the margins, and orogenic collapse resulting from the lithospheric delamination and upwelling of the asthenospheric materials within the interior of the orogenic belt

. Second episode Neogene 每 Quaternary of orogenic lithosphere root formation . Intracontinental orogeny and lateral growth of the orogenic belt (mainly in Miocene).

extrusion of the North China and the Yangtze blocks, denudation and planation. . Eastward fan-shaped opening of the Japan Sea eastwards and subduction of the South China Sea eastwards . Second episode of the upwelling of the mantle plume
. Westward developing

and crustal thickening and thinning accommodated contemporaneous geological events in eastern and western China. Taken together, this evidence reinforces the bidirectional indentation 每 extrusion model proposed previously and, as discussed below, is consistent with a 3-D model for collision-induced asthenospheric flow. 7.2.2.3. Continental root 每 plume tectonics. China and its environs may be viewed as a bell-shaped entity with an eastward &mouth* reflecting shortening in the west and extension in the east (Fig. 29). The addition of a depth parameter to 2-D models allows for developing and testing the &continental root 每plume tectonics* model. Because the Tibetan plateau orogen can be described in terms of orogenic lithospheric root formation and eastern Chinese rift zone in terms of plume-like mantle upwelling, &continental root 每 plume tectonics* is predicated on their common genetic association, the one compensating the other in 3D space (Deng et al., 1996a; Fig. 31). 7.2.3. The role of cratons Our proposed tectonic division of China (Fig. 4, Table 1) assumes that persistence of the Tarim, Max,

. Intracontinental

orogenic igneous activities in both the margins

continental rifting up to Baikal, cessation of both the Japan Sea and the South China Sea due to departure from the mantle plume resulted by the lithospheric drifting eastwards . Extensive basaltic magma eruption

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269

Fig. 31. Cartoon showing the plastic flow-off eastwards of the asthenospheric materials beneath the China continent (after Deng et al., 1996a). L〞lithospher; A〞asthenosphere.

Ordos, and Yangtze blocks as a single stable craton reflects their overall buoyancy and existence of their respective lithosphere roots (Deng et al., 1994e,f ). Throughout the Mesozoic and Cenozoic, this cohesive unit formed a stable barrier between eastern and western Chinese lithosphere, evidenced by the lack of significant igneous activity, deformation, and earthquakes, being independent of tectonic activity in the latter. The cratonic plexus defines northern and eastern margins of a compressional orogen, Tibet on the inner side, Cenozoic rifting to the outer side, being consistent with a stabilizing role for lithosphere roots (Section 3). He isotopic studies of the gases indicate a significant mantle contribution in the eastern rift basins and northern and eastern margins of the Tibet plateau, reflecting significant mantle degassing (Table 15), whereas He isotope compositions of the gases for the central cratonic group show unambiguous crustal character (Table 15) (Wang, 1989; Shen et al., 1991; cf. Deng et al., 1996a). In our view, this effect is fundamental to understanding the complex relationship between eastern and western China.

7.3. Subcontinental mantle dynamics 7.3.1. A three-dimensional mantle extrusion model Although two-dimensional planar models are able to explain extensional stress fields associated with South China Sea and Japan Sea opening, rapid asthenosphere upwelling is still required to account for oceanic crust formation. Because rapid thermal thinning of lithosphere, basalt magmatism, regional uplift, and crustal thickening in Tibet cannot be explained by two-dimensional models, we need to include depth (pressure) as a third dimension. According to Figs. 2, 3, 4, and 20 and the discussion in Section 6, plumelike mantle upwelling beneath eastern China likely drove from the 400-km-depth discontinuity as depicted in the profile in Fig. 31. Seismic tomographic images beneath China (Liu et al., 1989) show that the mantle is relatively cool between 600 and 800 km in depth and relatively hot between 220 and 400 km in depth, while at depths of c. 110 km, it is hot beneath eastern China and cool beneath western China. Given that western boundary conditions are relatively constrained, the eastern margin is characterized

Table 15 3 He/4He value (after Deng et al., 1996a) Crust
3 4 ?8

Mantle
?5

East China rift basin
?6

Tengchong
?6

Cratonic blocks Sichuan (1 每 2) ? 10 Crustal source
?8

Tarim (6 每 22) ? 10
?8

Junggar (7 每 54) ? 10
?8

Gansu (4 每 26) ? 10? 8

He/ He

10

10

(5 每 7) ? 10 (3 每 7) ? 10 30 每 70% of mantle contribution

270

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by Pacific plate subduction, and the 400 km discontinuity reflects the olivine 每 spinel mantle transition; these results support our projected model (Deng et al., 1996a; Fig. 31). Sub-Tibetan lithospheric roots are considered to have resulted from a combination of horizontal compression and downward mantle convergence, such that displaced asthenosphere extruded towards the relatively unconstrained eastern boundary. Subduction-related stress fields enhanced by lateral asthenospheric flow produce upwelling plumes leading to basalt magmatism and marginal basin formation. Implicit in this model is the dynamic relationship between Tibetan 每Himalayan orogenic compression, eastern Chinese continental rifting, and a conjugate asthenosphere flow field. Although differences in tectonic styles are considerable, they both are controlled by the same dynamic mantle system. A similar 3-D mantle extrusion model was also presented recently by Flower et al. (1998). 7.3.2. Summary In summary, we highlight key implications of the 2-D and 3-D models. (1) Bidirectional indentation in the west is responsible for inducing the eastward block extrusion and causing lithosphere extension in the east. (2) The formation of orogenic lithospheric roots in the west is responsible for inducing eastward flow of the asthenosphere, leading, in turn, to plumelike mantle upwelling. (3) Contrasting tectonic processes in the east and west are genetically related and are fundamentally complementary in character. Given the complexity of these systems, however, future research is needed to help distinguish natural geological effects from those produced by experimental simulations. These include the geometries of natural and experimental indentor surfaces, multiple-indentation events, drift associated with indentation fronts, the number and complexity of fault responses, heterogeneity of the indentor and deformed terrain, and the small but significant eastern boundary constraints.

8. Conclusions 1. Chinese continental lithosphere comprises three tectonic domains, i.e.: (1) eastern China, a region characterized by continental rifting, extensional basins, and basaltic volcanism; (2) central China, a

cratonic complex with low-heat flow (40 每50 mw/ m2), including the Tarim, Erdos, and Yangtze blocks, welded by pre-Cenozoic orogenic belts; and (3) western China, a region comprising the Qinghai 每Tibet 每 Himalaya orogen. 2. Relatively thin crust ( f 35 km) and lithosphere ( f 70 km) in eastern China is believed to reflect mantle upwelling, whereas near-normal crust ( f 45 km) and thickened lithosphere (>200 km) in central China suggests a mantle lithosphere root resembling those of Kaapvaal and Siberian cratons. Thickened crust ( f 70 km) and lithosphere (>150 km) in western China reflects the advanced development of an orogen root. 3. Relative buoyancy of the central Chinese lithosphere probably contributed to its long-term tectonic stability. However, gravitational instability produced by the thickening of denser western Chinese lithosphere and the resulting subsidence and eventual delamination of orogenic roots are believed to have produced postorogenic extensional collapse. 4. Although the geological relationships suggest compressional forces caused both orogen root development and the uplift of mountains and plateaus, extensional stress resulting from gravitational collapse is believed to have induced lithosphere and crustal thinning with significant reduction of topography. In contrast, the lower density of lithospheric mantle roots is likely to stabilize cratonic blocks in the asthenosphere. 5. Accordingly, the term &continental roots 每plume tectonics* has been adopted to describe the configuration and dynamic condition of subcontinental lithosphere and upper mantle beneath China, and it is proposed that supracrustal tectonic forms represent surface expressions of, and responses to, deep continental roots 每 plume tectonics*. 6. The prevailing view is that the western orogenic belt is not genetically related to eastern continental rifting. In contrast, such a relationship is inherent in the continental roots 每 plume tectonic model, such that formation of the orogenic lithosphere root forces eastward flow of the asthenosphere, leading to mantle upwelling beneath eastern China. 7. We suggest that continental roots 每 plume tectonics represent the sum of processes leading to the formation and evolution of the Eurasian supercontinent.

J.F. Deng et al. / Earth-Science Reviews 65 (2004) 223每275

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Acknowledgements Many thanks to Dr. Martin F.J. Flower for his invitation to write this paper, for subsequently constructive discussions, and reviewing and revising the manuscript, especially helping with improving in English. Thanks also to Drs. Paul Robinson and Nguyen Hoang for manuscript reviewing. All of the reviews substantially improved the manuscript. This work is supported by the National Natural Science Foundation of China (No. 40234048, No. 49973012, No. 40103003, No. 40172025, No. 49772107, No. 49772155, No. 49802005), the Science Foundation of the Ministry of Land and Resources of China (No. 20001010202, No. 200101020401, No. 9501101), the National Key Projects for Basic Research (No. G1998040807, No. 2001CB711002, No. 2002CB412600), China Geological Survey Project (No. 200113900018), and the &211* project &Mantle Materials and Deep Processes* from the China University of Geosciences, Beijing, as well as the IGCP-430. We thank Dr. Wang Yang for translation of Sections 6, 7, and 8, and acknowledge Miss Liu Cui and Mr. Xu Liquan for drafting figures.

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