Earth-Science Reviews 65 (2004) 223 – 275
www.elsevier.com/locate/earscirev
A new model for the dynamic evolution of Chinese
lithosphere: ‘continental roots–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 –
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 – Tibet – 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 – 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 – 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 – 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 – 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 – 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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J.F. Deng et al. / Earth-Science Reviews 65 (2004) 223–275
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.
2. Continental roots– 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 – 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 –
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,
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 – 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 – 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 – 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– T –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
J.F. Deng et al. / Earth-Science Reviews 65 (2004) 223–275
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 – 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 – chemical principles that govern phase
equilibria, kinetics, and element partitioning.
Here, we draw on Cenozoic and Mesozoic examples in developing the ‘continental roots – 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 –plume tectonics
2.3.1. The lithosphere– 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—
225
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 – asthenosphere boundary occurs at 60 –
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 – Tibet plateau,
and c. 104 km beneath its eastern margin, the ‘North –
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 – 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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J.F. Deng et al. / Earth-Science Reviews 65 (2004) 223–275
absent down to at least c. 460 km deep (8301 Workgroup, 1988). According to magnetotelluric studies,
lithosphere – asthenosphere boundary is located at 60 –
80 km in depth beneath the north China plain (Liu,
1985), 100– 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– 140 and 160 –200 km beneath the southern and northern parts of the Qinghai –
Tibet plateau, respectively (Wu et al., 1989), 140 –160
km beneath Qaidam –Qilianshan –Beishan (Zhu and
Hu, 1995), and 80 km beneath the North – 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 – asthenosphere boundary occurs at 60– 80 km in depth beneath
eastern China, no low-velocity zone beneath Yangtze
block (Song et al., 1992) and Tarim block, 116 –121
and 74 – 92 km beneath Qaidam and Qilianshan,
respectively (An et al., 1993; Zhuang et al., 1992),
120– 130 km beneath the Qinghai – Tibet plateau and
67 – 74 km beneath the North – South tectonic belt
(Zhuang et al., 1992), and 90 – 110 km beneath
Sanjiang – Bayan Har – Qaidam (Zhou et al., 1991).
Beghoul et al. (1993) suggested that the Qinghai –
Tibet lithosphere thickness was c. 205– 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– 186 km for Qinghai – Tibet, 190 km for
Tarim, 160 km for upper Yangtze, 72– 90 km for east
China, and 56 – 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– 80 km beneath the east Chinese rift zone (Deng et
al., 1990a), and c. 80 – 130 km beneath the northern
margin of the Qinghai –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– 100 km
for the Qinghai – Tibet collisional belt. Assuming that
the lithosphere –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 – Tibet,
referred to as ‘lithosphere roots’, and thinned lithosphere beneath east China corresponding to upwelling
mantle ‘plumes’. On this basis, ‘continental roots –
plume tectonics’ provides a generalized framework for
describing the Chinese continental lithosphere.
2.3.2. The crust – mantle boundary
All geophysical data defining the Moho depth
beneath China yield a consistent picture, giving average crustal thickness of 30 –35, 40 – 50 km, and c. 68
km for eastern China, central China, and the Qinghai –
Tibet plateau, respectively, varying gradationally between the first two regions and at margins of the
Qinghai –Tibet plateau (Feng, 1985). Surface elevations correspondingly vary from >4500 m above sea
level in the Qinghai – Tibet plateau, to 2000 – 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 –
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 – south tectonic belt at c. 102 – 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 –plume
J.F. Deng et al. / Earth-Science Reviews 65 (2004) 223–275
227
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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J.F. Deng et al. / Earth-Science Reviews 65 (2004) 223–275
Fig. 3. Cartoon of China lithosphere/crust contour. (a) The orogenic belt of Qinghai – 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 – Tibet orogenic belt. 2—Upper mantle isodensity
layer, density boundary (after Wang and Cheng, 1982); layer density: r = + 1, layer thickness: 150 – 850 m; layer density: r = 1, layer
thickness: 150 – >850 m; layer density: jrj = 1, layer thickness: 9 – 150 m.
J.F. Deng et al. / Earth-Science Reviews 65 (2004) 223–275
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 –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 –
229
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 –westtrending anomaly associated with the Qinghai – Tibet
plateau, a negative anomaly extending from Xinjiang
through Erdos to Sichuan – 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 – 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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J.F. Deng et al. / Earth-Science Reviews 65 (2004) 223–275
rium associated with recent redistribution of mantle
mass (Wang and Cheng, 1982; Ding, 1991). For
example, the Qinghai –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)
Crustal
thickness
(km)
Lithosphere
thickness
(km)
Upper mantle
density
(g/cm3)
Elevation (m)
Magmatism
Surface heat
flow
(mw/m2)
Earthquake
Qinghai – Tibet
intracontinental
orogenic belt
Central
China
cratonic
group
East China
continental
rift zone
70 – 80
(crustal root)
40 – 50
(normal)
30 – 35
(thinned)
150 (orogenic
lithosphere
root)
200
(cratonic
lithosphere
root)
3.10 – 3.22
70 (mantle
plume)
3.4 – 3.65
>4500
(plateau, high
mountain)
Intracontinental
shoshonitic
series, Ms/two
mica granites
Variable; north
part: 40 – 50;
margins:
70 – 90;
inner part:
100 – 300
Many and
strong
Recent
tectonics
Transpressional
Low-velocity
and highconductivity
layer in crust
Present
2000 – 1000
(basin and
plateau)
None
3.23 – 3.30
V 500 (basin
and range)
Continental
rifting basalts
40 – 50
60 – 70
None
(except for
interblock
boundaries)
Stable
(whole
uplifting
and
subsiding)
Absent
Many and
strong
Transtensional
Present
(1) the Qinghai – Tibet plateau west to Xining –
Chendu– Kunming and (2) the eastern China east to
Changchun – Taiyuan – Guangzhou, separated by a
lower density region beneath Xinjiang – Erdos –
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 – 3.65 g/cm3 (Qinghai – Tibet
plateau), 3.23 – 3.30 g/cm 3 (eastern China), and
3.10 – 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 –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.
3. Chinese lithosphere structure
Petrologic approaches to studying crust – 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 – tectonic events, crust – mantle relationships
J.F. Deng et al. / Earth-Science Reviews 65 (2004) 223–275
231
Table 2
Vp and Vs of minerals (km/s)
Vp
Vs
Quartz
Alkali feldspar
Plagioclase
(An = 24)
Plagioclase
(An = 53)
Biotite
Amphibole
Diopside
Garnet
Olivine
6.05
4.09
6.01
3.34
6.22
3.40
6.57
3.53
5.26
2.87
7.04
3.81
7.70
4.38
8.53
4.76
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 –
Tibet – 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 – Tibet –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 –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– 800 jC, 1.0–
1.1 GPa pressure (c. 35– 40 km depth), high-pressure
granulite forms in response to orthopyroxene breakdown, while at 400 – 800 jC, 1.2 – 1.6 GPa pressure
(c. 45 – 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 –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 –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 – 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)
Intermediate
granulite
(Opx 15;
Cpx 15;
Ga 5; Pl 65)
Intermediate
granulite
(Opx 10;
Cpx 10;
Bi 10; Ga 5;
Pl 65)
Acid
granulite
(Opx 2;
Cpx 3;
Ga 5;
Pl 60;
Qz 30)
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)
7.25
4.01
6.80
3.82
6.55
3.65
6.36
3.73
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
Granite
Amphibolite –
facies gneiss
Amphibolite
Intermediate
granulite
Basic
granulite
Acid
granulite
Basic
eclogite
Dunite
Vp (km/s)
Vs (km/s)
6.05 – 6.50
3.59 – 3.80
6.30 – 6.60
3.75
6.92 – 7.18
3.90
6.70 – 6.80
3.90
7.10 – 7.40
3.90 – 3.93
6.41 – 6.57
3.58 – 3.74
7.90 – 8.20
4.44 – 4.46
8.10 – 8.45
4.58 – 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– garnet transition (Wu et
al., 1994; Deng et al., 1997a,b). The asthenosphere
beneath both extensional rift zones and the Qinghai –
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 –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 –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) 223–275
Table 5
Crust – 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.
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 –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).
3.4. Tectonic implications
4. The record of magmatic activity
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
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 –plume hypothesis (Table
6). We note that lithologic assemblages in ‘welded’
J.F. Deng et al. / Earth-Science Reviews 65 (2004) 223–275
Table 6
Intracontinental igneous petrotectonic assemblages (after Deng et
al., 1996a)
Tectonic
setting
Petrotectonic
assemblage
Intracontinental (1) Oceanic
orogenic belt
closure and
continent – continent
collision orogeny:
blueschist and
eclogite; collisiontype PTt paths;
calc – alkaline
and high-K
calc – alkaline
magmatism;
(2) Intracontinental
orogeny:
intracontinental
subduction boundary:
biotite/two mica/
muscovite granites;
distributed shortening –
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
Roots – plume
tectonics type
Orogenic
lithosphere
root and
mountain root
De-rooting:
(1) Early stage:
lithosphere
thinning with
crustal root
(2) Late stage:
lithosphere
thinning without
crustal root
Mantle plume
Continental
lithosphere root
continent –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 – continent collision (Table 7;
Deng et al., 1996b). Because oceanic crust is consid-
235
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 – 41 Ma) and granitic
(55 – 41 Ma) rocks (Liu et al., 1990) range from calc –
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 –800 jC rather than the lowtemperature H2O-saturated melts (cf. 600– 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
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Table 7
Sequence of events: north Qilian Caledonian orogenic belt (after Deng et al., 1996a)
Seafloor spreading
and subduction
Oceanic closing and
continent – continent
collision
Intracontinental
subduction orogeny
Postorogenic
collapse and uplift
Cratonization
(a) MORB, Pt2, O1, O2, O3
(a) Blueschist (eclogite)
(subducted oceanic
crust; arc volcanic
rocks, pelitic rocks),
460 – 440 Ma (O3)
(b) Marine molasse, S1,
unconformable contact
with pre-Silurian
Two mica-granites,
417 – 404Ma(S2 – S3)
(a) A-type granites, D1, D2
(b) Retro-metamorphism
of blueschist,
400 – 380Ma (D1 – D2)
(c) Continental molasse, D1,
conformable contact
with pre-Devonian
Young platform, C
(b) OIB, Pt2, a2, O1, O2, O3
(c) Subduction arc volcanism,
a2, O1, O2, O3
(d) Subduction arc granites,
532 – 443 Ma (a2 – 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– 60 km, and Himalayan
crust having been 60– 75 km in thickness. Experimental petrologic data show that muscovite-bearing
melts are generated by partial melting of pelitic
compositions at depths of 20 –40 km. The Himalayan
muscovite/two-mica granites were formed in 7– 28
Ma, whereas retrograde metamorphism of the central
crystalline core occurred in 10 – 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.
J.F. Deng et al. / Earth-Science Reviews 65 (2004) 223–275
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 – 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
237
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 –
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 –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 – alkaline affinity.
Despite petrologic and geochemical indications of
volcanic rocks with island arc – continental margin
affinity, their emplacement over 1000 km north of
the MTB c. 20 –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 – 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 – 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 – 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 – Tibet shoshonites are also
relatively rich in TiO2 (1.04 – 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 –alkaline; (2) are characterized by relatively high
whole-rock HFSE contents; and (3) show Ti-rich
clinopyroxene compositions.
Qinghai – 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 –130 km (Lai et al.,
1996), being consistent with geophysical constraints
J.F. Deng et al. / Earth-Science Reviews 65 (2004) 223–275
239
Fig. 10. Cartoon showing orogenic lithospheric delamination and magma underplating (after Deng et al., 1995a).
on asthenosphere depth of c. 90– 120 km (An et al.,
1993) or c. 140 –160 km (Zhu et al., 2002). In contrast
to muscovite/two-mica granites, most Qinghai –Tibet
orogenic shoshonitic rocks, generated in the lower
parts of thickened continental crust (>50 – 60 km
depth), have no negative Eu anomalies. Average
crustal thickness in the northern Qinghai – Tibet orogenic belt is 60– 70 km with a high-velocity layer
(Vp = 7.3 –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 – 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 –
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 – Tibet – Himalaya orogen recently entered
a stage of orogenic collapse, reducing the lithosphere
thickness to c. 120 km from c. 150 – 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 –plume tectonics model (see Table 6).
A-type granitoids include the alkaline syenite –
quartz syenite series and also peralkaline granites,
and their volcanic equivalents include alkaline trachyte – 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 – 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 –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 – 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 – 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 – plume
tectonics model and continental dynamic processes
in general.
5. Orogen dynamics of the
Qinghai – Tibet–Himalaya
5.1. Paired igneous belts
Following the India –Asia ‘hard’ collision (c. 45
Ma), the Qinghai – Tibet –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– Yumen volcanics on the northern margin of
the Qinghai – 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– 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 –29 Ma with possible upper and
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241
Qinghai – Tibet plateau volcanic rocks occurred along
adjacent regions of the Altyn strike – 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 –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 – Tibet – Himalaya intracontinental orogenic igneous rocks (after Deng et al., 1994d,
1996a). Shoshonite series volcanic rocks belt: 1—northern Tibet;
3—Kekexili; 5—Yutian – 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 – 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 – Jinsajiang suture; BLS—Bangonghu – Lanchangjiang suture; YZS—Yaluzhangbujiang suture;
MCT—main central thrust fault; MBT—main boundary thrust fault.
lower limits of 38 –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 – 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 – Tibet plateau. Recent petrologic
and isotopic age studies (Deng, 1989) and detailed
mapping (Sun, 1992) show that they formed c. 14 – 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 – 23 Ma in age
(Deng et al., 1994a,b) and were paired with the Hoh
Xil volcanic zone. Since the Pleistocene (Q), northern
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 –Tibet – Himalaya orogen. Such episodicity is consistent with the results of the British– China Joint
Expedition Team of Tibet Plateau Geology (1990)
which yielded evidence on the basis of tectonic
deformation for three (45 –30, 30 –5, 5 –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 – south axis. The
first episode occurred in the Oligocene, with southern
and northern boundaries corresponding to the Yarlung – Zangbojiang and Tuotuohe – Jinsha Jiang
sutures (northern Qiangtang block), respectively. This
episode is considerd to be responsible for the formation of the Gangdese – 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 – Tibet – 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 – Tibet – Himalaya
orogen. However, lateral growth of the orogen is
hardly to be explained by the Indian subduction model
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but may be explained by the model of India – 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 – Beishan lithosphere has
not significantly underthrust the Qilian – Qaidam –
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 – Tibet – 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 –Zangbo section, and may be referred to as
‘Yarlung– Zangbo’ type. Following the ‘hard’ India –
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–
Zangbo types of crustal suture reflect intracontinental
block collision and continent – 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 – 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 – China Joint Expedition Team of Tibet Plateau Geology, 1990), crustal shortening in the regions north to the Qinghai –
Tibet plateau must also be considered. Average crustal
thickness of Tarim, Tianshan, and Altai are c. 48– 53,
50 –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 – 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.
J.F. Deng et al. / Earth-Science Reviews 65 (2004) 223–275
5.6. Mantle convergence and lithosphere root
formation
The India – 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-
243
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 – Tibet – 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 – 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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J.F. Deng et al. / Earth-Science Reviews 65 (2004) 223–275
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 – Tibet
plateau
Rapid uplift of the Qinghai – 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 – 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.
sodes of basalt magmatic activity: Paleogene and
Neogene –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 – Quaternary activity occured mainly in
smaller rifted basins and is concentrated in the following three zones from west to east: Daxinganling –
Taihangshan, Changbaishan – 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–
East Africa rift associations and is strongly contrary to
Yanshanian (J – K) calc – alkaline, HKCA, and shoshonitic character in eastern China. The bimodal character is dominated by mafic compositions with minor
felsic eruptives (trachyte – comendite) appearing only
in Tianchi of Changbaishan (Deng, 1988).
Magmas from rifts with larger extensional velocity
produced alkali basalt (Ne-norm = 0 – 5%), minor
basanite (Ne-norm = 5– 9%), and locally, hy-normative tholeiite. Rift flank eruptions formed mostly
basanite (Ne-norm z 9 – 10%), and, locally, nephelinite, with a lateral zonation resembling that of the East
African rift. The spatial variation of Neogene –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 –
Subei.
6.2. Crust –mantle thermal structure
6.1. Continental rift magmatism
The distribution and petrological characters of
volcanism in eastern China indicate two main epi-
Temperature distribution is a significant parameter
in discussing crust –mantle dynamics. Eastern Chinese Neogene –Quaternary basalts contain abundant
J.F. Deng et al. / Earth-Science Reviews 65 (2004) 223–275
245
Table 8
Average K2O and total alkali contents in basalts (%; after Deng, 1988)
Neogene – Quaternary
K2O
K2O + Na2O
Paleogene
Daxinganling – Taihangshan zone
Changbaishan – Tanlu fault zone
Southeast coast zone
West
flank
Axis
East
flank
West
flank
Axis
East
flank
West
flank
Axis
East
flank
1.59
1.19
5.10
5.04
Inland p coast
2.18
5.78
2.18
5.78
2.00
5.60
2.37
6.03
2.13
5.55
1.75
4.96
2.16
5.75
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 –mantle thermal
structure beneath eastern China in the context of rapid
Neogene– Quaternary lithosphere thinning. Normally,
the crust –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
North China
Bohai – Subei
0.80
1.22
3.47
3.87
Inland p coast
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 – 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 –350 jC in the depth
of 10 – 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 – 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 –
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.
J.F. Deng et al. / Earth-Science Reviews 65 (2004) 223–275
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 –
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 – 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
247
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 –
Quaternary, respectively; the latter includes historic
activity in regions such as Changbaishan and Wudalianchi. Thermal thinning induced by the Neogene –
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 –T equations for magmas
using available data on activities of the chemical
components involved. The cross point of several P –
T curves is the P –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– 100 km,
mostly between c. 50 and 80 km, i.e., 50 –100 km
(mainly 60 – 80 km) in northern and northeastern
China, 43 –56 km in Hannuoba, 55 – 72 km in northern Jiangsu, 50 km in the coastal region of southeastern China, 65 km along the Xinchang – Mingu zone,
and 100 km along Juxian– 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 – 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– 68 km (Song et al.,
1986), and S-wave velocities (Vs) indicate 60 – 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 –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
Depth to asthenosphere
Amount of
interstitial magma
Pliocene –
quaternary
Miocene
50 – 85 km,
(average 70 km)
65 – 100 km,
(average 85 km)
50 (Beijing) – 55 km
(Bohai, Northern Jiangsu),
(average 53 km)
3 – 10%,
(mainly 6 – 8%)
3 – 10%,
(mainly 5 – 7%)
15 (Beijing) – 18%
(Bohai, Northern
Jiangsu)
Paleogene
layer Vp values beneath the north China plain and
northern China are 7.2 – 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 –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 –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
J.F. Deng et al. / Earth-Science Reviews 65 (2004) 223–275
249
Fig. 15. Schematic profile of Paleogene basin – 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 –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 –16 km) and
the lower limit of listric faults (5– 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 – 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– 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– 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– 10 in the Paleogene when
large boundinage tectonics developed, changing to
103 – 104 in the Neogene– 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 – 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 – asthenosphere boundary; g1 and g2 are subcrustal lithosphere and crustal
viscosity, respectively.
(g1) needs to decrease to 1020 – 1022 Pas 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 – 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 – 1022 Pas, for depths < 75
km, and 1019 –1020Pas 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 –Quaternary
(characterized by doming tectonics). In other words,
the Paleogene mantle g1 value would be greater than
that in the Neogene – Quaternary. However, as the
Paleogene asthenosphere was shallower than that during the Neogene– Quaternary (Table 9), Paleogene g1
should be less than that in the Neogene –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– 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 – 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
Heat flow Example
anomaly
Extremely
rapid
thermal
thinning
Large
boudinage
tectonics
Large
basins;
(mantlederived
basalts)
Local
Paleogene
anomalies in eastern
China
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)
Two phase
uplift,
(1) main
uplift
accompanied
with thermal
thinning;
(2) uplifting
accompanied
with thermal
relaxation
ditto
ditto
One-phase
uplift—
terminated
by thermal
thinning
Regional Yanshanian
heat flow (Jura –
anomaly Cretaceous)
in eastern
China
Regional
anomalies
Neogene –
Quaternary
in eastern
China
J.F. Deng et al. / Earth-Science Reviews 65 (2004) 223–275
The case for slow-rate thermal thinning depicted in
Table 10 is still tentative and requires further study.
Widespread mantle-derived, crust – mantle-mixed, and
crustal source Yanshanian (J –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 –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– 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-
251
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– 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 –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 – 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 ¼ ðH qA Þ þ ðh qL Þ
h=H ¼ ðqL qA Þ=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 –140 km, respectively, by a comparison between
the Archean lithospheric root of 180 –220 or 200 km
and the Cenozoic lithosphere of 75 –80 or 60 –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– 3700 m a.s.l. (Ma, 1985) and those
for the Neogene – Quaternary at 1200 –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 – 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 – 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 – Quaternary alkali basalts
(Tables 9 and 11), estimates of qA in the Paleogene
and for the Neogene –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)
Paleogene
Neogene –
Quaternary
qmagma
(g/cm3)
F
qL
(g/cm3)
qA
(g/cm3)
h/km
H/km
2.82
2.77
0.15
0.07
3.33
3.33
3.25
3.29
2
0.7
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– 2000 km
and show temperatures c. 100– 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 –
Quaternary
rifting
83
157
58
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J.F. Deng et al. / Earth-Science Reviews 65 (2004) 223–275
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 – 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
253
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 –
North China region, Shuangliao –lower Liaohe River,
and Mudanjiang, respectively. In the Neogene– Quaternary, the primary plume also consisted of four
subplumes (V, VI, VII, and VIII) corresponding to
the Changbaishan, Zhangjiakou – 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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J.F. Deng et al. / Earth-Science Reviews 65 (2004) 223–275
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 – 3 GPa and
1300– 1500 jC, whereas hydrous basalts are generated at 2 –3 GPa and 1200 – 1400 jC (assuming melt
contains c. 1% H2O; Deng et al., 1992a,b,c,d).
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 – C –
H – 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).
Basalt eruption temperatures, calculated from meltmineral geothermometers to be c. 1130 –1200 jC,
allow interpolation of melt segregation temperatures
of about 1200 – 1250 jC at 2 – 3 GPa assuming adiabatic magma ascent (Deng et al., 1992a,b,c,d). Melt
segregation temperatures of about 1270– 1300 jC at
2 –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 –
C – H – 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– 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,
J.F. Deng et al. / Earth-Science Reviews 65 (2004) 223–275
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-
255
sphere – 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 –
Quaternary.
Fig. 21. Schematic section of the shape and structure of the mantle plume beneath northern China in (a) Paleogene and (b) Neogene – 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 – 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 –
north China plain, Shuangliao –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 – 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 – 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 – 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 – 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 –
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 – Quaternary
‘subplumes’ indicate that volcanism shifted 525 –
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 –Miocene peneplain surface. Prior to the inception of Baikal rifting, the Late
Cretaceous –Paleogene weathered crust persisted up
to the latest Oligocene – 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 –17 Ma) among a reasonably definitive range
of 116 K – 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-
J.F. Deng et al. / Earth-Science Reviews 65 (2004) 223–275
257
gene – 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 – 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. 22. Distribution of basalts in Hannuoba – Datong volcanic
basin. 1—N1 basalt; 2—N2 basalt; 3—Q1 basalt; 4—Q2 basalt (after
Deng, 1988).
Fig. 23. Distribution of basalts in Jingyu – Jingbohu vocanic basin
(after Deng, 1988). 1—N1 basalt; 2—N2 basalt; 3—Pleistocene
basalt.
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– Datong and Jiyu – 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 –Datong and Jinyu – 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 – 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.
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.
6.6.3. Volcanic ‘‘riveting’’ of the lithosphere?
The rate of lithosphere drift during active volcanism is apparently very low, only 0.04 – 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
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
Fig. 25. Asymmetrical spreading model (after Deng, 1988).
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.
J.F. Deng et al. / Earth-Science Reviews 65 (2004) 223–275
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 –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 – 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 –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,
259
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 –Pb and K – Ar ages of 253 – 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 – 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 – Lincang, (Yunnan Province)
extend southwards into Burma, Thailand, and Malaysia and show ages of 237– 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 –122 and c. 225– 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 – 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 – west Shangyi – Chifeng –
Longhua –Jianping – Beipiao in northernmost China.
The overthrust ‘Inner Mongolia axis’ has Triassic
muscovite/two-mica granites, and, to the south, the
foreland fold –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 – 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 – 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 –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 – Hingan – Zhangguangcai Hercyjan orogen.
Studies of ophiolite belts between Hegenshan in the
east to Souloushan– Karamery and west Junggar in
J.F. Deng et al. / Earth-Science Reviews 65 (2004) 223–275
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 – 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 –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 – 202 Ma (Heilongjiang;
Hong et al., 1991), 218.8 and 235.9 Ma (Fanshan and
Yangyuan), 234 Ma (Liyuanhekanzi), 240– 190 Ma
(Saima), 224 Ma (Antuqinglizi; Zou et al., 1988),
185.8 and 203 Ma (Ulongshan, Fengning, and Guangtoushan– Pngquan; Wang et al., 1994), and 248.8–
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 – 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 – continent’
collisions. The ultrahigh pressure coesite eclogites in
Dabieshan and Sulu have the Sm –Nd isochron ages
of 244– 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 –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 – continent collisional oroge-
261
nies. Assuming, therefore, that continental subduction
was the dominant process, horizontal shortening and
strike – 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 – Ar ages of 200 –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 –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 –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 –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 – 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 –171 Ma (Wang et al., 1994; Ye et
al., 1991). This extends northeastwards and connects
with the Yanji pluton belt (208 – 170 Ma; Zhou et al.,
1992) and so-called gneiss– granite series (206 –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 – 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– 106.1 Ma (cf. Deng et
al., 1996a). These granites record continent –continent
or arc – continent collisions following the closure of
the Mesozoic Mongolia – Okhotsk – 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 – 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 –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 –1.5 and 1.4 –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 – 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 – 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 – 1000, 400 –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– 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 – 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.
J.F. Deng et al. / Earth-Science Reviews 65 (2004) 223–275
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 – 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 – 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 –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-
263
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 –Baikal
arc, Irkusk arc, Mongolia arc, and Qilianshan –
Luliang arc (Liu, 1980). Liu (1980) referred to the
trace of the tip of these arcs as the ‘north– south belt
of East Asia’ (Helan –Kangdian line), pointing out its
correspondence to the Angora – 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 –
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 –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 –
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 –
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 –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 – west compressional
field and, in turn, north – south striking Riyueshan in
east Qinghai and the inland Qinghai Lake (cf. Deng,
J.F. Deng et al. / Earth-Science Reviews 65 (2004) 223–275
265
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 – 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 – Quaternary. Fault: M.O.F.—Mongolia – 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)
Paleogene
(E)
Main events
in the southern
part of East
Asia related
to the India
indentation
Main events
in the northern
part of East
Asia related
to the Siberia
indentation
. The Red
. The right-
. F1
lateral strike –
slip fault of
the north
margin of the
North China.
. The
. B1
The
Sundaland
Mongolia –
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 –
lateral strike –
slip fault of the
slip fault.
north margin
of the Qinling –
Dabie (?).
. The left-lateral . F1
The Red
river rightstrike – slip fault
lateral strike –
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 –
slip fault.
.
.
Neogene(N) – .
Quaternary
(Q)
.
.
.
E – Q (?)
.
Source
.
In comparion
to the model
of the
indentation
experiments
(Fig. 28)
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.
J.F. Deng et al. / Earth-Science Reviews 65 (2004) 223–275
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 – Altyn
Tagh – 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
267
events since the collision of Tibet were taken from the
discussion in Section 5. The timing of late Oligocene – 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– 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 –Hoh Xil region was covered by Pliocene – Pleistocene volcanic rocks, whereas later studies suggested that the volcanic rocks were Miocene
(24 – 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 –
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 –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 –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 –27 Ma) and
early Miocene (23 –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,
268
J.F. Deng et al. / Earth-Science Reviews 65 (2004) 223–275
Table 14
Roughly comparison of the geological events on the east and west
China and adjacent areas (after Deng et al., 1996a)
Qinghai – Tibet –
Himalaya orogenic
belt
. First episode of the
Paleogene
(Oligocene) formation of orogenic
lithospheric root
. Intracontinental
orogeny
. Intracontinental
orogenic igneous
activities in both
the margins
. Uplifting,
Latest
Oligocene – denudation and
earliest
planation of the
Miocene
orogenic belt
. Second episode
Neogene –
Quaternary of orogenic
lithosphere root
formation
. Intracontinental
orogeny and
lateral growth
of the orogenic
belt (mainly in
Miocene).
. Intracontinental
orogenic igneous
activities in both
the margins
East China and
adjacent areas
. First episode of the
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 –
northeast
China block
. Widespread
extensive basaltic
magma eruptions
Table 14 (continued )
Qinghai – Tibet –
Himalaya orogenic
belt
East China and
adjacent areas
. Third episode
. Since Pleistocene,
Neogene –
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
. Rapid eastward
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
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
and crustal thickening and thinning accommodated
contemporaneous geological events in eastern and
western China. Taken together, this evidence reinforces the bidirectional indentation – 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 – 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 –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 –
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,
J.F. Deng et al. / Earth-Science Reviews 65 (2004) 223–275
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
He/ He
10
8
Mantle
5
10
East China
rift basin
Tengchong
Cratonic blocks
Sichuan
6
6
(5 – 7) 10
(3 – 7) 10
30 – 70% of
mantle contribution
Tarim
8
(1 – 2) 10
Crustal source
Junggar
8
(6 – 22) 10
Gansu
8
(7 – 54) 10
(4 – 26) 10 8
270
J.F. Deng et al. / Earth-Science Reviews 65 (2004) 223–275
by Pacific plate subduction, and the 400 km discontinuity reflects the olivine – 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 –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 –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 –Tibet –
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 –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 – 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 – 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 – 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) 223–275
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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