Accretionary Tectonics of the Western Kunlun Orogen, China: A
Paleozoic–Early Mesozoic, Long-Lived Active Continental Margin
with Implications for the Growth of Southern Eurasia
W. J. Xiao, B. F. Windley,1 D. Y. Liu,2 P. Jian,2 C. Z. Liu, C. Yuan,3 and M. Sun4
State Key Lab of Lithospheric Evolution, Institute of Geology and Geophysics, Chinese
Academy of Sciences, Beijing 100029, China
(e-mail: [email protected])
ABSTRACT
Our new SHRIMP U-Pb zircon ages from the Western Kunlun Orogen allow us to constrain the history of an active
continental margin developed on the southern boundary of the Tarim block from the Ordovician to the Triassic. A
492 Ⳳ 7-Ma dacite from Yixieke extrusive rocks that contain 220 Ⳳ 5 -Ma reheated zircons is interpreted as an intraoceanic arc complex that accreted to the Tarim block. The Yirba granodiorite has a continental arc geochemical
signature, a 471 Ⳳ 5-Ma U-Pb crystallization age, and 491 Ⳳ 3 -Ma inherited zircons. It formed during the first, early
Paleozoic stage of an active continental margin arc that was juxtaposed to the south against the Kudi high-grade
gneiss complex, the Buziwan ophiolite, and the Yixieke volcanic and sedimentary rocks. Zircons from a paragneiss
in the Kudi gneiss complex range in age from 398 Ⳳ 12 to 1345 Ⳳ 31 Ma; the oldest reflect protolith ages of a gneissic
continental block (incorporated into the trench), and the youngest may represent the age of a refoliated high-grade
fabric created during accretion. The Buziwan ophiolite occupies a thrust sheet tectonically overlying the Kudi gneiss
complex. A leuco-gabbro pegmatite, with a zircon age of 403 Ⳳ 7 Ma and ca. 490-Ma inherited zircons, and the North
Kudi granite, with a zircon age of 408 Ⳳ 7 Ma, were emplaced during the second mid-Paleozoic stage of the active
continental margin. The Akarz subduction-related granite that has a 214 Ⳳ 1 -Ma zircon crystallization age formed
during the final, early Mesozoic stage of the active margin. The long-lasting active continental margin in the western
Kunlun forms a key, well-documented section of the Andean-type margin that extends from the Caucasus to the
Qinling.
Online enhancements: color versions of figures 3 and 4.
Introduction
the Tarim block to the north and the Tethyan domain to the south, and it sheds light on the tectonic
architecture of the Tibetan plateau immediately to
its south. However, the Paleozoic tectonic evolution of the WKO has been contentious (see Jiang et
al. 1992; Yao and Hsü 1994; Matte et al. 1996; Mattern et al. 1996; Pan 1996; Yang et al. 1996; Yuan
1999; Mattern and Schneider 2000; Xiao et al.
2002a, 2002b, 2003a; Yuan et al. 2002a, 2002b) because of the lack of reliable isotopic ages of certain
key tectonic units, in particular the Kudi ophiolite,
which is situated in a possibly middle Paleozoic
(Akaz) suture (Matte et al. 1996; Mattern and
Schneider 2000; Pan 1996; Sobel and Arnaud 1999;
Cowgill et al. 2003) that extends eastward to the
Lapeiquan suture in the Altyn–East Kunlun (fig. 1).
The Western Kunlun Orogen (WKO), located along
the northern periphery of the Tibetan plateau, is a
1000-km-long mountain belt extending from the
Pamir syntaxis in the west to the Altyn–East Kunlun Orogen in the east (fig. 1). Its Paleozoic to early
Mesozoic orogenic history is of considerable importance for the reconstruction of paleo-Asia because it occupies a key tectonic position between
Manuscript received June 3, 2004; accepted January 28, 2005.
1
Department of Geology, University of Leicester, Leicester
LE1 7RH, United Kingdom.
2
SHRIMP Laboratory Beijing, Institute of Geology, Chinese
Academy of Geological Sciences, Beijing 100037, China.
3
Guangzhou Institute of Geochemistry, Chinese Academy
of Sciences, Guangdong, Guangzhou 510640, China.
4
Department of Earth Sciences, University of Hong Kong,
Hong Kong SAR, China.
[The Journal of Geology, 2005, volume 113, p. 687–705] 䉷 2005 by The University of Chicago. All rights reserved. 0022-1376/2005/11306-0005$15.00
687
688
W. J . X I A O E T A L .
Figure 1. Schematic map of the Kunlun-Qinling ranges and adjacent regions showing the Paleozoic–early Mesozoic
active continental margin, marked by crosses (modified after Sobel and Arnaud 1999; Xiao et al. 2002b; Cowgill et
al. 2003; Roger et al. 2003); F p fault. Area of figure 2 is indicated by the box.
This suture is a remnant of the paleo-Tethyan
ocean in central Asia (Pan 1996; Sobel and Arnaud
1999). Knowledge of the isotopic ages of the Kudi
ophiolite and spatially associated granites is essential to understand the tectonic evolution of paleoTethys (Pan 1996; Zhou and Graham 1996; Wang
1997; Sobel and Arnaud 1999; Yue and Liou 1999;
Yuan et al. 2002a, 2002b).
In this study, we report new zircon U-Pb SHRIMP
ages from the Kudi ophiolite and two key granites
that have well-established structural age relationships with the ophiolitic rocks, in order to document specific stages in the crustal evolution of the
WKO. We have integrated all these SHRIMP zircon
ages with our recent structural, geochemical, petrological, geochronological, and tectonic data from
the WKO (Yuan 1999; Xiao et al. 2002a, 2002b,
2003a; Yuan et al. 2002a, 2002b). We use published
geochronological (Xu et al. 1994, 1996; Bi et al.
1999; Sobel and Arnaud 1999; Cowgill et al. 2003,
2004a, 2004b; Gehrels et al. 2003a, 2003b) and geochemical data (Deng 1995; Zhang et al. 1996; Yuan
1999; Jiang et al. 2002; Wang et al. 2002; Yuan et
al. 2002a, 2002b, 2004) to reevaluate the origin of
the plutonic and volcanic rocks and their interre-
lationships within a supra-subduction zone setting.
Significant advances in understanding the WKO
have been made in the last several years, including
improvements in understanding of the timing and
patterns of deformation (Zhou et al. 2000; Xiao et
al. 2002a, 2002b, 2003a), the origin of the granitic
plutons (Yuan et al. 2002a, 2002b), the nature of
the Ordovician arc (Sobel and Arnaud 1999; Cowgill et al. 2003), and the accretionary tectonics in
the southern part of the WKO (Xiao et al. 2002a,
2002b, 2003a). This article summarizes the broad
geological environments and structures of the
WKO and presents a new model to explain the major tectonic events within the context of the accretionary orogens of southern central Asia.
Regional Geology
The WKO is divisible into the North Kunlun, South
Kunlun, and Tianshuihai domains, separated by the
Akaz and Mazar-Kangxiwar faults, respectively (fig.
2). The North Kunlun domain represents the basement of the Tarim block, and the South Kunlun
domain is mainly composed of various tectonic assemblages, including the Buziwan (Kudi) ophiolite,
Journal of Geology
KUNLUN ACCRETIONARY TECTONICS
689
Figure 2. Schematic tectonic map of the Kunlun ranges and adjacent regions showing the position of the Kudi area
and major pluton age distribution (modified after Xu et al. 1996; Sobel and Arnaud 1999; Xiao et al. 2002b; Cowgill
et al. 2003; Gehrels et al. 2003a, 2003b). Area of figure 3 is marked by a box. MKF p Mazar-Kangxiwar fault.
the Yixieke arc, the Kudi gneiss, and the Xiananqiao arc (Xiao et al. 2002a, 2002b, 2003a), the main
subjects of this article. The Tianshuihai domain is
mainly a huge accretionary wedge that records a
late Paleozoic–early Mesozoic subduction-related
orogenic process (Xiao et al. 2002a, 2003a). For
more detailed stratigraphy and structures, readers
are referred to Matte et al. (1996), Pan (1996), Mattern and Schneider (2000), and Wittlinger et al.
(2004). Figure 3 shows the main rock units, described below mainly in time sequence. In this article, we use the new geological time scale (Gradstein et al. 2004) to analyze tectonic stages or
geological events.
The northernmost unit, near Akaz Daban (pass;
fig. 3), comprises the basement of the Tarim block
(Matte et al. 1996; Mattern and Schneider 2000; Pan
1996), which is composed of Proterozoic gneisses,
schists, migmatites, stromatolite-bearing limestones, clastics, and cherts, overlain by Sinian conglomerates, tillites, clastics, and carbonates (fig. 3).
On the southern side of the gneisses, a fault (figs.
3, 4) marks the southern boundary of exposed rocks
of the Tarim block. South of the fault, greenschists
form 20- to 1100-m-thick layers within an 800-m–
1.5-km-thick succession (the Sailajaz Group of
Yuan et al. 2004) of stromatolite- and crinoidbearing Neoproterozoic to Early Cambrian bedded
marbles that are cut by 30-cm-thick basic dikes (of
unknown age) and contain layers of slate at least
100 m thick. Yuan et al. (2004) show that the greenschists are large-ion lithophile and light rare earth
element (LREE) enriched, have relatively high Th/
Nb and La/Nb ratios, and within-plate character-
istics, together with a wide range of Zr/P2O5 ratios
and a concomitant increase in Th/Nb ratios, suggesting crustal contamination. We interpret the
Sailajaz belt as a whole as a remnant of a continental rift shelf in which shelf limestones are cut
by rift dikes and are interbedded with mudstones
and rift lavas. The belt is situated on the southern
margin of the Tarim block and represents a halfgraben-shelf succession that bordered an ocean to
the south. The Sailajaz Group rests unconformably
on the Middle Proterozoic Tarim basement rocks
that consist of metamorphosed clastic, carbonate,
and volcanic rocks intruded by a 2.2-Ga granitic
pluton (Pan 1996; Yuan et al. 2004); this confirms
a continental basement beneath the shelf rift
succession.
South of the Sailajaz belt (fig. 3), Zhang (1997)
reported Ordovician crinoids in marbles, which, together with greenschists, were interpreted to be
fragments of a volcanic seamount in an accretionary wedge (Xiao et al. 2002b). This means that there
must have been ocean floor on which the seamount
was built, and therefore we suggest that the southern side of the Sailajaz belt marks a new position
for a suture zone along the Akaz fault (fig. 3). Closure of the ocean resulted in the Akaz suture, which
separates continental and continental-margin rocks
of the Tarim craton (North Kunlun domain) to the
north from accreted oceanic-derived ophiolitic and
arc rocks of the orogen to the south (South Kunlun
domain; fig. 3).
On the southern side of the accretionary wedge
is the Yirba arc-related, lineated, and foliated granodiorite (fig. 3), which has U/Pb zircon ages of
Figure 3. Tectonic map of the Kudi area, Western Kunlun Orogen (based on our field data, incorporated with those
of Matte et al. 1996; Mattern and Schneider 2000; XBGMR 1993; and Yin and Bian 1995). A color version of this
figure is available in the online edition of the Journal of Geology.
Journal of Geology
KUNLUN ACCRETIONARY TECTONICS
691
Figure 4. Cross section A–A along the line shown in figure 3. Key as in figure 3, except as indicated. KP p
knick point where section direction changes. See text for discussion. A color version of this figure is available in the
online edition of the Journal of Geology.
491 Ⳳ 3 and 471 Ⳳ 5 Ma (Yuan et al. 2002a, 2002b).
This forms the older active margin developed along
the southern margin of the Tarim block (Xiao et al.
2002b).
Kudi Arc-Ophiolite-Accretionary
Wedge Assemblage
An important assemblage of the South Kunlun domain was formerly known as the “Kudi ophiolite
suite” (Matte et al. 1996; Pan 1996; Mattern and
Schneider 2000) and is mainly composed of ultramafic rocks and volcanic and volcaniclastic rocks
and forearc sediments of the former Yishak Group
(Xiao et al. 2002a). The ultramafic rocks include
a southward-thrusted ultramafic-gabbroic klippe
about 3 km long and 1.5 km wide in the Buziwan
valley and 2–3-m-thick tectonic slices now imbricated within the Kudi gneiss in unnamed valleys
south of Kudi. The volcanic and volcaniclastic
rocks and forearc sediments crop out mainly to the
north of the ultramafic-gabbroic rocks, with a very
good section in the Yixieke Valley (fig. 3). Because
there are various rock types with different structures and geochemical signatures in this former
ophiolite suite, we use the term “Buziwan (or Kudi)
ophiolite” of Xiao et al. (2002a, 2003a) and Wang
et al. (2001, 2002) to encompass ultramafic and gabbroic rocks in the Buziwan valley and “Yixieke
forearc” for volcanic and volcaniclastic rocks in the
Yixieke valley. Accordingly, we describe the main
components of the accreted ophiolite-arc rock assemblage under these locality names.
The Yixieke extrusive rocks consist of massive
and pillowed basalts, boninites, tuffs, welded andesitic breccias and agglomerates, and calc-alkaline
lavas intruded by uncommon dolerite dikes; there
are no sheeted dikes (Matte et al. 1996; Pan 1996;
Sobel and Arnaud 1999; Yuan 1999). Three groups
of tholeiitic lavas were recognized by Wang et al.
(2002). Group 1 basalts have LREE-enriched,
chondrite-normalized REE patterns and Cr-Y values typical of island arc tholeiites and La/Sm-TiO2
ratios similar to those of the Mariana arc. Group 2
basalts have low K contents, marked negative Nb
anomalies, flat to slightly LREE-depleted REE patterns typical of transitional midocean ridge basalt
(T-MORB), Cr-Y values akin to those of island arc
basalts, La/Sm-TiO2 ratios comparable to those of
the Lau back-arc basin, and a supra-subduction
zone signature in Hf/3-Th-Nb/16 space. Buziwan
gabbros and diabase dikes that transect extrusive
rocks have geochemical signatures similar to those
of the group 2 tholeiites. Our dacitic sample 123
occurs within the group 2 basalts near the bottom
of the Yixieke volcanic pile. Group 3 basalts are
characterized by high Cr and low Y indicative of a
relatively high degree of partial melting derived
from a depleted mantle source, have low TiO2
(0.16–0.38 wt%), high Mg (Mg# p 62–72), normal
midocean ridge basalt (N-MORB)–normalized trace
element patterns, La/Sm-TiO2 ratios similar to
those of forearc boninites in the Izu-Bonin-Mariana
arc, and U-shaped REE patterns that are typical of
many boninites (Yuan 1999). The three groups of
Yixieke basalts are overlain by calc-alkaline lavas
that include basaltic andesites, andesites, and volcaniclastic rocks such as tuffs, welded andesitic
breccias, and agglomerates.
However, a recent detailed field and geochemical
study along the Yixieke valley has provided a new
tectonic framework in which five units are recog-
692
W. J . X I A O E T A L .
nized in the central part of the Yixieke extrusive
rocks (Yuan et al. 2005). The lowest, unit A, has
N-MORB-like geochemical characteristics, and the
overlying unit B suggests an enriched midocean
ridge basalt (E-MORB) affinity. The geochemistry
of overlying units C and D reflects the involvement
of a slab-derived component, possibly produced by
partial melting of a mantle source modified by
melt-rock interaction during upwelling of E-MORB
mantle. The uppermost unit, E, shows geochemical
features that can be explained by mixing of a MORB
component with melts from subducted sediments.
It is noteworthy that tholeiitic basalts have initial
143
Nd/144Nd and 87Sr/86Sr isotopic ratios ranging
from 0.5122 to 0.5123 (␧ Nd p 5.8–8.0) and from
0.7037 to 0.7050, respectively. Boninitic lavas are
characterized by high Al2O3/TiO2 values of 120, low
TiO2 and Al2O3 values, high SiO2 and Na2O values,
LREE-enriched patterns ((La/Yb)N p 1.5–2.0), and
␧Nd values lower than 3.0. These ratios and the distribution of major- and trace-element data point to
an origin in an incipient oceanic arc created by possible mixing of fertile oceanic island basalt, depleted subarc mantle, and fluids derived from a subducted slab (Yuan 1999). The geochemical data
show that the rocks are akin to evolved boninites
of the Mariana forearc (Yuan 1999; Wang et al. 2002;
Xiao et al. 2002b). Such boninites are typical of a
supra-subduction zone environment, where magma
generation is strongly influenced by aqueous fluids
(Hickey and Frey 1982).
The volcanic rocks are overlain by a 1500-mthick succession of Yixieke turbidites (fig. 3) that
includes ophiolite-derived debris flows, tuffaceous
and andesitic sandstones, and radiolarian cherts
(Wang 1983; Pan 1996). The Yixieke turbidites are
subdivided into lower and upper parts, with a tectonic contact between them (Wang 1983; Jiang et
al. 1992; Mattern and Schneider 2000; Xiao et al.
2002b). A maximum age for the sedimentation is
indicated by Late Ordovician–Silurian radiolaria in
the lowermost turbidites. Petrochemical data of the
lower turbidites suggest an origin in a forearc basin
(Fang 1998; Fang et al. 1998); this is consistent with
the predominantly forearc nature of the underlying
volcanic rocks.
Above the Late Ordovician–Silurian lower turbidites is a thrust, above which are the younger
upper turbidites that contain radiolaria of Late
Devonian–Early Carboniferous (Fang 1998; Zhou et
al. 2000) and possibly Carboniferous-Permian age
(Mattern and Schneider 2000). The thrust sheet is
composed of imbricated turbidites with secondorder thrust faults (Mattern and Schneider 2000;
Xiao et al. 2002b, 2003a). We interpret these post-
thrust turbidites as clastic debris deposited in a
forearc basin that underwent late thrusting.
The Buziwan ultramafic-gabbroic rocks make up
a 3-km-thick slab that contains sheared basal serpentinite, layered/foliated chromite-bearing dunite, harzburgite, clinopyroxenite, and gabbro. The
main body of unserpentinized dunite contains layers of clinopyroxenite and hornblendite. Harzburgites and dunites are traversed by veins of olivineorthopyroxene, clinopyroxenite, and asbestos
(Wang et al. 2001, 2002), and gabbros are cut by
dikes of gabbro (Jiang et al. 1992). The meaning of
a whole-rock mineral Sm-Nd isochron age of
651 Ⳳ 53 Ma on dunite, harzburgite, gabbro, and
plagioclase from the gabbro is uncertain (Ding et
al. 1996). The intrusive Akarz early Mesozoic
granodiorite (zircon age of 212–213 Ma; Yuan et al.
2002a, 2002b) contains three lenses up to 100 m
wide of dunite and gabbro (fig. 4). The main ultramafic slab has been thrust southward over the Kudi
gneiss complex (fig. 4).
Xiao et al. (2002b, 2003a) interpreted the Buziwan ophiolite as a substrate of Buziwan ocean
floor overlain by the Ordovician-Silurian suprasubduction Yixieke arc, in turn overlain by a Late
Ordovician–Silurian turbiditic forearc basin. However, the isotopic age of the ophiolite is not known,
and this uncertainty has led to contrasting speculations on its age: late Neoproterozoic (Wang 1983),
Proterozoic–early Paleozoic (Matte et al. 1996; Pan
1996), and late Paleozoic (Jiang et al. 1992; Yao and
Hsü 1994; Yang et al. 1996; Yin and Harrison 2000).
The Kudi gneiss complex (figs. 1–3) forms the
main ridge of the WKO. It is composed of hornblende/biotite gneisses that contain minor lenses
of schist, marble, phyllite, quartzite, and amphibolite that are cut by two generations of discordant
amphibolite dikes, which in the second generation
are subhorizontal and undeformed. The gneisses
contain a 10-m2 body of anorthosite (Zhou et al.
2001). The gneiss complex has been interpreted as
a Proterozoic microcontinent derived from the Tarim block (XBGMR 1993; Ding et al. 1996; Pan
1996) and as a metamorphosed Paleozoic accretionsubduction complex (Şengör and Okurogullari
1991; Şengör and Natal’in 1996; Zhou et al. 2000).
The 40Ar/39Ar dates on hornblende (452 Ⳳ 5 Ma)
and biotite (428 Ⳳ 2 Ma) from the gneisses, in relation to kinematic indicators, suggest that they
were affected by Late Ordovician–Early Silurian local ductile shearing (Matte et al. 1996; Zhou et al.
2000).
The North Kudi granite (Matte et al. 1996; Mattern and Schneider 2000; Jiang et al. 2002; Yuan et
al. 2002a) intrudes the Kudi gneiss complex (Zhou
Journal of Geology
KUNLUN ACCRETIONARY TECTONICS
et al. 2000) and now is in fault contact with the
Yixieke volcanic rocks (fig. 3). It has high d18O
(11.6%) and high 87Sr/86Sr ratios (ISr p
0.7097–0.7119) but slightly lower ␧Nd(t) values
(⫺3.8 to 1.4; Jiang et al. 2002) than the volcanic
rocks from the Yixieke arc, which have ␧Nd(t) values
from 1.4 to 4.4 and 87Sr/86Sr ratios ISr p
0.7054–0.7069 (Deng 1995). These relations suggest
that the source region of the granite involved subducted oceanic crustal sediment. The granite has
U/Pb zircon ages of 380.0 ⫹1.9/⫺0.7 Ma (Xu et al.
1994) and 405 Ⳳ 2 Ma (Yuan 1999; Yuan et al.
2002b). To get a more precise age and test the two
different ages, we conducted a new analysis using
SHRIMP zircon dating.
The Buziwan main dunite body and the Kudi
gneiss complex were stitched by the Akarz granitic
pluton, which has 40Ar/39Ar ages of 180 Ⳳ 10 and
221 Ⳳ 6.6 Ma (Xu et al. 1994) and a zircon U-Pb
age of 214 Ⳳ 1 Ma (Yuan et al. 2002a, 2002b).
Analytical Procedures
Locations from which the samples were collected
are shown in figure 3. Zircons were separated using
conventional heavy-liquid and magnetic techniques. Representative zircons were hand-picked
and, together with several examples of standard zircon TEM from the Research School of Earth Sciences (RSES), Australian National University,
mounted in epoxy resin and sectioned approximately in half, and the mount surfaces were polished to expose the grain interiors and then gold
coated.
Zircons were analyzed at the Chinese Geological
Academy of Sciences using SHRIMP II. The
SHRIMP data have been reduced according to the
method of Williams and Claesson (1987), Williams
(1992), Williams et al. (1996), Compston et al.
(1984, 1992), and Huang et al. (2004). Interelement
fractionation was estimated relative to the RSES
standard zircon TEM (417 Ma). The U, Th, and Pb
concentrations were determined relative to those
measured in the standard zircon SL13, which has
a U concentration of 238 ppm and an age of 572
Ma (Claoué-Long et al. 1995). Corrections for common Pb were made using the measured 204Pb/206Pb
ratios. Because of the small amount of 207Pb formed
in young (i.e., !1000 Ma) zircons, which results in
low count rates and high analytical uncertainties,
the determination of the ages for young zircons has
to be based primarily on their 206Pb/238U ratios
(Compston et al. 1992). Uncertainties in the isotopic ratios and ages in the data table (table 1) and
in the error ellipses in the plotted data are reported
693
at a 1j level, but the final ages on pooled data sets
are all 206Pb/238U ages reported as weighted means
at 95% confidence level. All age calculations and
statistical assessments of the data have been made
with the geochronological statistical software
packages ISOPLT/EX (version 2.00) and SQUID 1.0
of Ludwig (1999, 2001).
SHRIMP U-Pb Geochronology
Yixieke Volcanic Rocks (Sample 123). We sampled
this dacite from a 5–10-m-thick flow within basaltic lavas at the southernmost end of the main body
of extrusive rocks, which is located 50 m south of
the bridge just north of the North Kudi pluton, in
order to constrain the maximum age of the ophiolitic lavas and to establish their isotopic age relationship with the Buziwan ultramafic-gabbroic
rocks. This is one of five dacitic flows near the base
of the extrusives. In the sample, plagioclase dominates over alkali feldspar in a 2 : 1 ratio, which is
typical of dacites.
Fifteen analyses of zircons yielded U/Pb ages,
with a number of analyses showing concordance
(fig. 5a). The major analyses plot as a group straddling the concordia and give a weighted mean 206Pb/
238
U age of 492 Ⳳ 9 Ma (n p 8, MSWD p 0.96), as
these data are located along or near the concordia
and show little variance. These zircons are mainly
clear and elongate with bipyramidal terminations,
and some are fragments of grains with pyramidal
terminations. A smaller group gives a weighted
mean 206Pb/238U age of 220 Ⳳ 5 Ma (n p 5,
MSWD p 1.16; fig. 5a; table 1). These grains are
equant and have prismatic magmatic shapes. Because these data are located along or near the concordia and have similar error ellipses (fig. 5a; table
1), we accept the latter as a concordant date.
Buziwan Pegmatite (Sample 120). We sampled this
hornblende leuco-gabbro pegmatite dike in Buziwan Valley (fig. 3) below the chromite mine (fig.
4), where gabbros are cut by gabbro dikes (Jiang et
al. 1992). Thirteen analyses of zircons yielded U/
Pb ages, with a number of analyses showing concordance, as these data are located along or near
the concordia and show little variance (fig. 5b).
Most analyses plot as a group straddling the concordia and give a weighted mean 206Pb/238U age of
403 Ⳳ 7 Ma (n p 11, MSWD p 0.47). The zircons
that yield the 403 Ma age are clean, thin, and long,
and nearly all are prismatic; we interpret 403 Ma
as the age of formation of the pegmatite dike. Two
analyses (11.1 and 12.1 in table 1) are statistical
outliers from this group and have an age of a ca.
490 Ma; these zircons are equant, with pyramidal
Table 1.
Summary of U-Th-Pb SHRIMP data on zircons from the Kudi ophiolite and associated granites
Grain,
U
Th
spot
(ppm) (ppm)
Sample
1.1
2.1
3.1
4.1
5.1
6.1
7.1
8.1
9.1
10.1
11.1
12.1
13.1
14.1
15.1
Sample
1.1
2.1
3.1
4.1
5.1
6.1
7.1
8.1
9.1
10.1
11.1
12.1
13.1
Sample
1.1
2.1
3.1
4.1
5.1
6.1
7.1
8.1
9.1
10.1
11.1
12.1
13.1
14.1
15.1
16.1
Sample
1.1
2.1
3.1
4.1
5.1
6.1
7.1
8.1
9.1
10.1
11.1
12.1
232
Th/
U
238
Pb∗
(ppm)
206
Pb∗/206Pb∗
207
%
Pbc
206
Pb/238U
ages (Ma)
206
Pb∗/235U
207
Pb∗/238U
206
Value Error (%) Value Error (%) Value Error (%)
119, Arkarz pluton (36⬚48.258⬘N, 76⬚56.351⬘E):
547
298
.56
16.1
1.10
215 Ⳳ 6
.0620
6.0
.290
499
286
.59
14.9
1.67
216 Ⳳ 6
.0696
14
.327
495
250
.52
14.3
.98
211 Ⳳ 6
.0710
13
.326
769
353
.47
23.3
.60
222 Ⳳ 6
.0551
6.4
.266
312
155
.51
9.15 3.81
209 Ⳳ 6
.0600
17
.271
489
298
.63
14.6
1.66
217 Ⳳ 6
.0555
9.6
.262
410
178
.45
12.2
2.28
215 Ⳳ 6
.0449
16
.210
552
298
.56
16.1
.87
213 Ⳳ 6
.0482
12
.223
418
152
.38
12.3
2.69
212 Ⳳ 7
.0540
20
.248
666
143
.22
19.2
1.27
210 Ⳳ 6
.0523
9.3
.239
463
150
.33
13.4
2.24
209 Ⳳ 6
.0505
9.4
.229
339
159
.48
10.4
2.09
222 Ⳳ 6
.0540
16
.260
622
320
.53
17.0
.87
200 Ⳳ 5
.0480
5.2
.208
2013 1367
.70
60.3
.40
220 Ⳳ 6
.0500
2.6
.239
276
184
.69
14.4
2.17
371 Ⳳ 10 .0541
9.6
.442
120, Buziwan leuco-gabbro pegmatite (36⬚48.105⬘N, 76⬚56.721⬘E):
186
275
1.53
10.2
1.41
391 Ⳳ 11 .0540
14
.466
276
229
.86
15.5
1.27
403 Ⳳ 11 .0513
6.9
.456
173
117
.70
9.83 2.32
403 Ⳳ 11 .0556
16
.495
227
134
.61
12.9
1.73
407 Ⳳ 11 .0587
6.9
.527
407
264
.67
23.3
.73
413 Ⳳ 11 .0552
4.2
.504
1030
200
.20
57.4
.29
404 Ⳳ 10 .0530
2.2
.473
968
187
.20
53.7
.30
402 Ⳳ 10 .0539
6.7
.478
411
574
1.44
23.1
.98
404 Ⳳ 11 .0551
4.5
.491
212
139
.68
12.3
1.89
412 Ⳳ 11 .0523
8.5
.476
155
213
1.42
8.77 2.35
403 Ⳳ 11 .0535
11
.476
292
171
.60
20.1
1.49
488 Ⳳ 12 .0583
5.0
.632
298
169
.59
20.6
.95
494 Ⳳ 13 .0590
3.9
.648
77
140
1.89
4.40 7.45
387 Ⳳ 12 .0570
25
.490
121, Kudi biotite gneiss (36⬚48.188⬘N, 76⬚59.751⬘E):
537
115
.22
69.3
.31
900 Ⳳ 22 .0789
1.2
1.630
484
98
.21
57.9
.65
836 Ⳳ 20 .0751
1.6
1.433
937
126
.14
93.7
.35
707 Ⳳ 18 .0720
1.1
1.151
482
171
.37
79.0
.41 1,120 Ⳳ 26 .0842
1.5
2.204
532
122
.24
62.0
.46
818 Ⳳ 20 .0787
1.5
1.468
601
46
.08
50.3
.41
597 Ⳳ 15 .0680
2.1
.910
99
110
1.15
5.76 6.08
398 Ⳳ 12 .0620
19
.550
430
155
.37
86.1
.44 1,345 Ⳳ 31 .1013
1.3
3.239
1022
27
.03
72.9
.25
513 Ⳳ 13 .0604
1.5
.690
687
66
.10
69.2
.65
711 Ⳳ 17 .0800
2.3
1.285
642
28
.04
50.4
.61
560 Ⳳ 14 .0672
3.2
.842
500
101
.21
61.8
.50
863 Ⳳ 21 .0783
1.6
1.546
765
89
.12
90.4
.51
827 Ⳳ 21 .0869
2.4
1.641
950
8
.01
65.6
2.16
488 Ⳳ 12 .0642
4.2
.696
594
25
.04
48.9
1.44
581 Ⳳ 14 .0645
4.9
.839
1151
116
.10 111
.24
684 Ⳳ 17 .0693
1.2
1.069
122, biotite granite of the North Kudi pluton (36⬚53.905⬘N, 76⬚58.909⬘E):
53
116
2.25
3.05 9.66
376 Ⳳ 15 .0410
61
.34
105
90
.88
5.68 1.77
385 Ⳳ 13 .0570
18
.487
233
435
1.93
12.9
1.28
397 Ⳳ 10 .0551
6.0
.483
131
193
1.52
7.43 3.29
400 Ⳳ 11 .0456
18
.402
282
167
.61
16.1
1.10
409 Ⳳ 10 .0539
6.0
.488
718
425
.61
42.0
.44
423 Ⳳ 11 .0545
2.5
.509
162
147
.94
9.64 1.48
424 Ⳳ 11 .0525
10.0
.493
335
390
1.20
19.9
1.72
423 Ⳳ 11 .0520
5.6
.486
427
260
.63
23.7
1.12
400 Ⳳ 10 .0558
5.9
.492
400
241
.62
22.8
.98
410 Ⳳ 10 .0539
4.1
.488
96
110
1.19
5.50 2.99
403 Ⳳ 11 .0551
15
.491
411
282
.71
23.0
.80
404 Ⳳ 10 .0538
5.0
.479
6.6
14
13
6.9
17
9.9
16
12
20
9.7
9.8
16
5.9
3.7
10.0
.0339
.0341
.0333
.0351
.0329
.0342
.0339
.0336
.0330
.0331
.0329
.0350
.0315
.0347
.0593
2.6
2.7
2.7
2.6
2.8
2.7
2.8
2.7
3.4
2.7
2.7
2.9
2.6
2.6
2.7
15
7.4
16
7.4
5.0
3.4
7.2
5.2
8.9
11
5.7
4.7
25
.0625
.0646
.0646
.0651
.0662
.0646
.0644
.0646
.0660
.0645
.0787
.0797
.0618
2.8
2.7
2.9
2.7
2.7
2.6
2.6
2.7
2.7
2.8
2.6
2.6
3.3
2.9
3.1
2.8
3.0
3.0
3.3
19
2.9
3.0
3.4
4.2
3.0
3.6
4.9
5.6
2.8
.1498
.1384
.1160
.1898
.1353
.0971
.0637
.2320
.0829
.1165
.0908
.1432
.1369
.0786
.0944
.1120
2.6
2.6
2.6
2.6
2.6
2.6
3.1
2.6
2.6
2.6
2.7
2.6
2.6
2.6
2.6
2.6
62
18
6.6
18
6.6
3.6
10
6.2
6.4
4.9
15
5.6
.0600
.0616
.0635
.0639
.0656
.0677
.0680
.0679
.0640
.0656
.0646
.0646
4.0
3.5
2.7
2.9
2.6
2.6
2.8
2.6
2.6
2.6
2.9
2.6
Journal of Geology
Table 1
KUNLUN ACCRETIONARY TECTONICS
695
(Continued)
Grain,
U
Th
spot
(ppm) (ppm)
232
Th/
U
238
Pb∗
(ppm)
206
Pb∗/206Pb∗
207
%
Pbc
206
Pb/238U
ages (Ma)
206
Pb∗/235U
207
Pb∗/238U
206
Value Error (%) Value Error (%) Value Error (%)
Sample 123, dacite of the Yixieke calc-alkaline rocks (36⬚57.502⬘N, 76⬚59.005⬘E):
1.1
657
204
.32
1.78 1.78
477 Ⳳ 12 .0558
2.9
.586
2.1
596
103
.18
.67
.67
491 Ⳳ 12 .0579
1.6
.630
3.1
640
112
.18
.62
.62
481 Ⳳ 12 .585
1.9
.623
4.1
995
208
.22
69.9
1.00
502 Ⳳ 12 .0656
1.9
.732
5.1
685
115
.17
46.9
.98
490 Ⳳ 12 .0626
3.6
.681
6.1
580
174
.31
41.9
1.54
512 Ⳳ 13 .0670
7.3
.763
7.1
538
226
.43
16.2
1.56
219 Ⳳ 6
.0528
6.2
.251
8.1
877
627
.74
26.5
1.54
220 Ⳳ 6
.0519
4.8
.248
9.1
578
183
.33
17.5
1.62
220 Ⳳ 6
.0424
9.2
.203
10.1 1225
192
.16
85.9
.73
502 Ⳳ 12 .0582
2.2
.650
11.1
475
68
.15
32.1
1.27
482 Ⳳ 12 .0579
4.1
.621
12.1 1405
313
.23
44.5
1.91
229 Ⳳ 6
.0511
5.5
.255
13.1
487
273
.58
14.3
2.03
212 Ⳳ 6
.0481
8.5
.222
3.9
3.0
3.2
3.2
4.4
7.7
6.7
5.5
9.6
3.4
4.9
6.1
8.9
.0761
.0789
.0772
.0809
.0789
.0827
.0345
.0346
.0348
.0810
.0777
.0362
.0334
2.6
2.6
2.6
2.6
2.6
2.6
2.6
2.6
2.7
2.6
2.6
2.6
2.7
Note. Errors are 1j. Pbc and Pb∗ indicate the common and radiogenic portions, respectively. The error in standard calibration was
0.66%. Common Pb percentage was corrected using measured 204Pb.
terminations. We consider these to be inherited
from the main host gabbro.
Kudi Gneiss Complex (Sample 121). We selected
this biotite gneiss, which comes from just 300 m
south of Kudi (fig. 3), in order to date the gneiss
complex. Zircons show variable shapes; some are
euhedral and long, others are broken and short, but
the majority are prismatic. Sixteen analyses of zircons yielded U/Pb ages that all show severe discordance (fig. 5c), suggesting extensive Pb loss. As
indicated in table 1, the ages range from 398 Ⳳ
12 to 1345 Ⳳ 31 Ma (for interpretation, see below).
North Kudi Pluton (Sample 122). We sampled the
North Kudi pluton (fig. 3) in order to compare its
SHRIMP age with published single-zircon U-Pb
ages of the pluton that range from 380 to 405 Ma.
A biotite granite sample yielded a single population
of zircons that are short, idiomorphic, prismatic,
and clear. The SHRIMP analyses (table 1; fig. 5d)
produced a 206Pb/238U age of 408 Ⳳ 7 Ma (n p 11,
MSWD p 1.17) that is identical to the U/Pb zircon
age of 404.0 Ⳳ 3.1 Ma of Yuan (1999) but significantly older than the U/Pb zircon age of 380.0⫹1.9
⫺0.7
Ma of Xu et al. (1994).
Akarz Pluton (Sample 119). We sampled this biotite granite (fig. 3) where it intruded the main Buziwan dunite body, leaving lenses of chromitelayered dunite (at the chromite mine) in the granite
(fig. 4), in order to document its age and to provide
a more precise upper age limit on the convergent
tectonic processes. Zircons are transparent, euhedral, prismatic, needle-like crystals that have magmatic morphologies. Fifteen analyses of zircons
yielded U/Pb ages, with a number of analyses showing concordance (fig. 5e). Most analyses plot as a
group straddling the Concordia and give a weighted
mean 206Pb/238U age of 213 Ⳳ 3 Ma (n p 14,
MSWD p 1.17). One analysis (15.1 in table 1) is
statistically an outlier from this group and indicates some inheritance from a ca. 371-Ma source.
Interpretation
Buziwan Ultramafic-Gabbroic Rocks and Yixieke Lavas.
The Buziwan ultramafic-gabbroic rocks in
the Kudi ophiolite form an important indicator of
an oceanic basin that was eliminated during the
accretionary process along the southern boundary
of the Tarim block. The unserpentinized Buziwan
dunites contain layers of hornblendite in what we
regard as a metamorphic foliation fabric. The predominant dunite-harzburgite composition of the
ultramafic rocks suggests that they are metamorphic restites after high degrees of partial melting of
a lherzolite mantle (Coleman 1977). This is consistent with the fact that pyroxenes in the harzburgites have low Al2O3 and TiO2 contents, suggesting that these rocks are residual mantle
peridotites after a high degree of partial melting
(Wang et al. 2002). Within a cogenetic magma sequence, restites of dunite-harzburgite composition
are complementary to hornblende gabbros, low-Ti
tholeiites, and high-Mg boninites (Beccaluva et al.
1983). All these rocks are different accreted components within an accretion-subduction complex,
although geochemistry suggests that they possibly
all belong to the same cogenetic magmatic sequence (cf. Khain et al. 2002).
The composition of chromian spinel is diagnostic
of particular tectonic environments (Dick and Bullen 1984). Wang et al. (2002) showed that Cr spinels
Figure 5. Concordia plot of SHRIMP U-Pb data for zircons from (a) sample 123, dacite from the Yixieke arc; (b)
sample 120, pegmatite from the Buziwan gabbro; (c) sample 121, biotite gneiss from the Kudi gneiss complex; (d)
sample 122, biotite granite of the North Kudi pluton; and (e) sample 119, biotite granite of the Akarz pluton.
Journal of Geology
KUNLUN ACCRETIONARY TECTONICS
in the Buziwan dunites have a Cr number (Cr# p
100Cr/(Cr ⫹ Al) p 60–67) that is typical of arcrelated ultramafic rocks associated with a subduction zone.
Although Wang et al. (2002) were unsure about
some age relationships of the basalts, they found
that the boninites are locally interbedded with and
overlie the group 2 basalts. In their interpretation
of the petrogenetic sequence, these authors placed
emphasis on the back-arc signature of the group 2
basalts, suggested that continued upwelling in the
back-arc allowed hydrous fluids from the subducting slab to trigger remelting of depleted refractory
mantle, so forming the boninites, and that group 1
island arc tholeiites enriched in LREE were created
last by renewed subduction, which permitted hydrous fluids and/or melts from the subducting slab
to interact with mantle rocks. They invoked mantle diapirism to initiate the back-arc spreading, following the idea of Karig (1971), and so discounted
the more modern alternative of trench rollback as
a mechanism to create the back-arc extension.
However, considering the documented geochemical characteristics in relation to current ideas on
arc magmatism, we suggest the following evolutionary scenario.
The most widely accepted model for the formation of back-arc basins and of supra-subduction
zone ophiolites depends on hinge rollback (Maruyama 1997; Shervais 2001). The several phases of
development of the Kudi ophiolite are precisely predicted by the model of Shervais (2001). The initial
phase of ophiolite formation in the forearc gives
rise to low-K, LREE-depleted tholeiites, which
range in composition from basalt to basaltic andesite and even dacite and often have flat REE patterns and trace elements that resemble MORB
(group 2 gabbros and tholeiites) and form by melting of MORB source asthenosphere before any introduction of fluids from the subducting plate. Our
dated dacite (sample 123) comes from these lowermost volcanic rocks, and thus its age provides a
reasonable estimate of the time of initiation of
ophiolite formation.
There are several alternative interpretations for
the origin of sample 123 dacite from near the base
of the Yixieke volcanics: (1) the later date (220 Ⳳ
5 Ma) is the result of Pb loss during a Triassic thermal event; (2) the dacite is a hypabyssal intrusion
related to Triassic magmatism, and the 492 Ⳳ 9Ma-dated zircons are xenocrystic grains; (3)
492 Ⳳ 9 Ma is the age of crystallization of the dacite, and the 220-Ma age is due to reheating by a
Triassic thermal event. Because the 220 Ⳳ 5-Ma
date is concordant, as indicated above, we exclude
697
the first possibility. If we treat the dacite as a hypabyssal intrusion related to Triassic magmatism
and the 492 Ⳳ 9-Ma-dated zircons as xenocrystic
grains, it is hard to reconcile the fact that a similar
Triassic granite (sample 119) and a Devonian granite (sample 122), which both occur nearby, do not
have any old xenocrystic grains. In addition, in the
Triassic rock assemblages no dacite has ever been
reported in the Kudi area. A close association of the
dacite with the Yixieke arc volcanic rocks leads us
to accept the last interpretation. This is in accord
with the fact that most early Paleozoic components, i.e., the leuco-gabbroic pegmatite and the
Yirba pluton, are closely related to the Kudi arcophiolite assemblage, which all have ca. 490-Ma
ages. A more detailed investigation is needed to test
the interpretation that we favor here.
The second phase generates tholeiites that are
low in Ti, enriched in LREE, and depleted in Nb
and Mg-enriched boninites. These melts are
brought about by an increasing flux of fluids from
the subducting slab. At this stage of the Shervais
(2001) model, an even higher fluid flux would give
rise to tholeiites even more enriched in LREE and
Ti (Shervais 2001), like the group 1 basalts.
The third and final phase in the formation of an
accreted active margin in relation to the suprasubduction ophiolite is the generation of calcalkaline basaltic andesites, andesites, and rhyolites,
together with hornblende-bearing diorites, tonalites, and trondhjemites. The Yirba hornblende
granodiorite belongs to this mature phase of development. Here 471 Ⳳ 5 Ma may be reasonably
regarded as the age of crystallization (Yuan et al.
2002a, 2002b), and we interpret 491 Ⳳ 3-Ma zircons as xenocrysts inherited from the early group
2 lavas. The 492-Ma dacite in the lowermost lavas
is close to the start of formation of the active margin, and 471 Ma (the age of latest granodiorite pluton) is close to the mature stage of its development.
We conclude that the Kudi ophiolite went
through a typical supra-subduction development
and has a corresponding geochemical signature, in
which case there is no need to invoke a single backarc model or a second-stage subduction in a backarc basin, as proposed by Wang et al. (2002). We
note that many of the stratigraphic and geochemical characteristics of the Kudi ophiolite are similar
to those of the Late Jurassic ophiolite in Hokkaido,
Japan, which Takashima et al. (2002) concluded
formed in a forearc rift basin above a suprasubduction zone; viz., harzburgites, pyroxenites,
and dunites containing chromian spinels with arcrelated chemistry, tholeiitic basalts with back-arc
basin-like chemistry interbedded with and overlain
698
W. J . X I A O E T A L .
by boninitic high-Mg andesites, calc-alkaline andesites, and turbidites in a forearc basin. Similar
relationships occur in the 1020-Ma Dunzhugur
supra-subduction ophiolite in Siberia (Khain et al.
2002).
The Kudi Gneiss Complex. A paragneiss from this
complex contains zircons that have a wide range
of dates, from 398 Ⳳ 12 to 1345 Ⳳ 31 Ma. The latest date, 398 Ⳳ 12 Ma, is located along the concordia, which indicates a concordant age. This
398 Ⳳ 12-Ma concordant date is an indication of
newly formed magmatic zircon in the Kudi complex; thus, a possible Early Devonian magmatic
event may predate an early phase of accretion partly
represented by the paragneiss. This is in good agreement with the similar ages of the North Kudi pluton and the Buziwan leuco-gabbroic pegmatite. In
addition, Matte et al. (1996) reported Ar-Ar ages of
380–350 Ma, and Zhou (1998) and Zhou et al. (2000)
reported Ar-Ar ages of 452–428 Ma in different segments of the Kudi gneiss complex, which they interpreted as records of metamorphism.
The other, earlier dates in the Kudi gneiss are
obviously for detrital zircons. Z. Hui (pers. comm.)
has obtained a U-Pb zircon date with a lower intercept age of 533 Ⳳ 21 Ma and an upper intercept
age of 1251 Ⳳ 23 Ma on a Kudi gneiss. The intermediate dates represent original ages modified by
extensive Pb loss caused by the youngest metamorphism. This is consistent with the fact that
granites intruding the Kudi gneisses yield Nd
model ages of 1.1–1.5 Ga (Yuan et al. 2002a). We
therefore interpret the earlier dates to reflect the
protolith ages of the continental gneissic block that
docked into the trench and the 452–350-Ma dates
as the time of peak metamorphism that gave rise
to the high-grade mineral assemblage and refoliated
fabric of the gneiss created during accretion.
Discussion
The History of the Kunlun Active Continental Margin. The tectonic history of the WKO has been
summarized by many workers from their available
geochemical, stratigraphic, structural, and tectonic
data, the most recent of which were Xiao et al.
(2002a, 2002b, 2003a) and Yuan et al. (2002a, 2002b,
2004). Our SHRIMP zircon dates provide new constraints on the timing of several key tectonic events,
which correspondingly require new interpretations.
Below we present a revised Paleozoic–early Mesozoic tectonic history of the WKO, illustrated in figure 6.
A passive continental margin rift succession existed before the early Paleozoic on the southern
border of the Tarim block, with an ocean (protoTethys) to the south (Pan 1996). As we discussed
earlier (Xiao et al. 2002b, 2003a), there was a period
of southward subduction of the passive marginal
sequence of the Tarim block in the Late Cambrian
to earliest Ordovician (fig. 6a). From the Early Ordovician, the ca. 490-Ma Yixieke arc–Kudi ophiolite complex accreted to the Tarim block, a northward subduction followed beneath the composite
Tarim accretionary margin, and thus the early stage
of an Andean-type magmatic arc developed on the
southern margin of the Tarim block (fig. 6b). This
is consistent with and explains the fact that the
dacite from the Yixieke arc volcanic rocks has a
formation age of 492 Ⳳ 9 Ma and that ca. 490-Ma
inherited zircons have been found in both the
403 Ⳳ 7-Ma leuco-gabbro pegmatite from the Buziwan valley and the 471 Ⳳ 5-Ma granodiorite from
the Yirba pluton. The intrusion of the Yirba granodiorite at 471 Ⳳ 5 Ma (Yuan et al. 2002b) and of a
similar 460⫹2.4
⫺2.5-Ma pluton nearby (Xu et al. 1996)
marked the mature phase of development of this
Andean-type active margin.
The mid-Proterozoic Kudi continental block was
approaching the subduction zone in the Late
Ordovician–Silurian to Early Devonian when the
ductile shear zone in the gneisses was created (fig.
6c). The Kudi gneiss complex underwent a relatively long accretionary process, as it contains obviously different tectonic components, including a
ca. 398-Ma paragneiss. In the Early to Middle Devonian, the ultramafic rocks of the ophiolite were
thrust over the Kudi gneiss complex, probably in a
trench. The collision and accretion at the leading
edge of the active margin assisted the creation of
a thrust-thickened mountain belt. Melting of underlying metasomatized mantle wedge created the
405-Ma lamprophyres (Zhou and Li 2000; fig. 6c).
Further development of the active margin generated the 408–380-Ma North Kudi granite and the
403-Ma leuco-gabbro pegmatite dikes that retain
inherited zircons from the early history of the margin. In the Late Devonian to Early Carboniferous,
a forearc basin created over the accretionary belt
received turbidites that contain clastic debris from
the eroding mountains (fig. 6d).
The resumption of northward-dipping subduction along the southern margin of the Tarim block
in the Permian to early Mesozoic was contemporaneous with the collision between the SiberiaAltaid continent and the northern margin of the
Tarim block (Heubeck 2001; Roger et al. 2003; Xiao
et al. 2003b). The 214-Ma Akarz subduction-related
granodiorite represents the third and final stage in
the development of the Andean-type margin (figs.
Figure 6. Sequential diagram showing the Palaeozoic–early Mesozoic tectonic evolution of the WKO. (a) Late
Cambrian to Early Ordovician; (b) Middle Ordovician; (c) Late Ordovician to Middle Devonian; (d) Late Devonian to
Early Carboniferous; (e) Late Carboniferous–Permian to early Mesozoic. See text for discussion.
700
W. J . X I A O E T A L .
Figure 7. Histogram with cumulative probability of all
dates in this study.
6e, 7). The southern Tarim active margin collided
with the Qiangtang block to the south in the Late
Jurassic (Roger et al. 2003). This collision terminated all subduction-related tectonic processes in
the northern WKO.
The Long-Lived Active Margin of Southern Eurasia.
From the distribution of the ages of granitic rocks,
a lack of magmatic activity from 350 to 220 Ma
(Yuan 1999) was previously interpreted as being related to cessation of subduction (Yuan et al. 2002a,
2002b). An analysis of a histogram with cumulative
probability also shows that there was such a gap
(fig. 7). However, the paleogeographic reconstructions of Nie et al. (1990) suggested northward subduction under the southern margin of the Tarim
block in the Early Permian, and this idea was supported by the discovery of Permian subductionrelated volcanic rock along the southern margin of
this subduction system (Matte et al. 1996; Mattern
et al. 1996; Pan 1996). Li et al. (1995) and Bi et al.
(1999) summarized Early Permian magmatic activity in the WKO. In the southern part of the South
Kunlun, Middle Devonian to Early Permian granodiorites were reported (Li et al. 1995; Xu et al. 1996;
Bi et al. 1999). Forearc accretion south of this possible late Paleozoic magmatic arc took place in the
Late Carboniferous to early Mesozoic (Xiao et al.
2002b, 2003a; Schwab et al. 2004).
In the meantime, in the eastern Kunlun, plutons
and volcanic rocks of arc affinity likely formed in
the 370–320-Ma interval (Dewey et al. 1988;
Schwab et al. 2004); in the Pamirs to the west, volcanism related to an arc began at ∼370 and 320 Ma
and most likely continued into the Triassic
(Schwab et al. 2004). Therefore, we propose that
during the late Paleozoic to early Mesozoic, the
southern active margin of the Tarim block still existed (Xiao et al. 2002a) and that there was only a
relatively short cessation of magmatic activity in
the Carboniferous. This kind of magmatic cessation is not uncommon in active margins, such as
the present-day western North American active
margin that is characterized by transform tectonics
with large-scale strike-slip faults and related basins
(Dickinson 1995).
This would be consistent with paleomagnetic
data that indicate that the Tarim block was moving
northward as a united plate in the Devonian to Late
Carboniferous (Li 1990; Yin and Nie 1996) and that
a collision occurred between the Tien Shan and
southern Siberia in Carboniferous-Permian times
(Windley et al. 1990). The period of northward drift
of the Tarim block resulted in a reduction in the
rate of subduction-accretion during the time period,
with a possible cessation in the Early Carboniferous, which would explain the relatively small
amount of documented igneous and accretionary
activity during this time. The WKO was submerged
and overlain by a thick pile of marine deposits in
the Late Devonian to Early Carboniferous (XBGMR
1993). Several early Paleozoic fossiliferous blocks
along this accretionary complex (Yin and Bian 1995;
Mattern et al. 1996; Pan 1996; Xiao et al. 2002b,
2003a) were probably incorporated by arc-parallel
strike-slip faulting (or associated extension) during
this time.
Therefore, we conclude, on the basis of the
SHRIMP dates and our other cited evidence, that
the southern Tarim active margin underwent an
early magmatic event at ca. 490 Ma that was followed by accretionary processes at ca. 400 Ma and
was finally terminated by Andean-type accretionary orogenesis at ca. 214 Ma, as indicated by the
histogram of the SHRIMP dates conducted in this
study (fig. 7) and other chronological data summarized by Pan (1996), Yuan et al. (2002a, 2002b),
Cowgill et al. (2003), and Gehrels et al. (2003a,
2003b).
In recent years, forearc accretion has gained increasing popularity as a process to explain the evolution of many orogenic belts, such as those in central Asia, the Arabian-Nubian Shield (e.g., Şengör
and Natal’in 1996), the Lachlan Orogen of eastern
Australia (Gray 1997; Gray and Foster 1998; Foster
and Gray 2000), and the Proterozoic Yavapai Orogen south of the Wyoming craton (e.g., Hoffman
1988). Based on the Japanese model, accretionary
orogens evolve largely by processes of forearc accretion (Isozaki et al. 1990; Şengör and Okurogullari 1991; Windley 1992; Şengör et al. 1993; Xiao
Journal of Geology
KUNLUN ACCRETIONARY TECTONICS
et al. 2003b). Our geochronological data and tectonic interpretation all indicate that the Paleozoic
to early Mesozoic WKO was a long-lasting, complicated accretionary orogen, because an early Mesozoic Andean-type active continental margin developed on the Paleozoic-accreted margin of the
Tarim block. This long-lived active continental
margin is characterized by a major accretionary
complex and forearc basin on its southern side.
This new information will shed light on an important controversy about the evolution of this part
of Asia, viz., whether the WKO was a collisional
orogen that resulted from either the collision of
various terranes between the Tarim and Qiangtang
blocks (Dewey et al. 1988; Jiang et al. 1992; Matte
et al. 1996; Mattern et al. 1996) or collapse of backarc basins (Yao and Hsü 1994) or an accretionary
orogen (Şengör and Okurogullari 1991; Mattern and
Schneider 2000; Xiao et al. 2002a, 2002b, 2003a).
From a more regional perspective, Sobel and Arnaud (1999) and Cowgill et al. (2003) compared the
WKO with the East Kunlun, and their work supports a general tectonic model for the northern Tibetan plateau in which an intermediate island arc
or composite terrane was accreted to the Tarim
block along a southward-dipping subduction zone.
Gehrels et al. (2003a, 2003b) summarized different
models for the formation of the northern Tibetan
plateau; they excluded a back-arc model but sup-
701
ported a general tectonic model like that proposed
here. The general multiple accretionary framework
of our tectonic model for the WKO is consistent
with all previous investigations in which accretionary tectonics played a key role together with
southward subduction followed by arc accretion,
subduction flip (northward subduction), and formation of an Andean-type active margin. Recent
tectonic analysis of accretionary complexes in
other areas of the Tibetan Plateau (Kapp et al. 2000,
2003a, 2003b; Yin and Harrison 2000; Aitchison et
al. 2001) have greatly increased our understanding
of the role of accretion on the southern side of the
WKO in this segment of southern Eurasia.
Our proposed tectonic scenario of a long-lived
active continental margin is comparable to that in
the East Kunlun and Qinling (figs. 1, 8; Molnar et
al. 1987; He et al. 1999; Sobel and Arnaud 1999;
Xiao et al. 2002a; Cowgill et al. 2003; Roger et al.
2003; Bian et al. 2004; Schwab et al. 2004), although
detailed aspects of the evolution could be different
(Gehrels et al. 2003a, 2003b) because of possible
orogen-parallel variations. There are also similarities farther east in the Qinling-Dabie Orogen,
where the North and South China blocks collided
by the Late Permian (Nie et al. 1990), with the most
active deformation period in the Late Triassic to
Early Jurassic (Zhao and Coe 1987; Enkin et al.
Figure 8. Permian to Early Triassic paleogeography of Eurasia, with emphasis on the continental blocks and related
orogenic belts in central-east Asia (modified after Heubeck 2001; Xiao et al. 2003b). Dark gray shows some Precambrian
blocks in the east.
702
W. J . X I A O E T A L .
1992; Gilder et al. 1999; Roger et al. 2003; Bian et
al. 2004; Schwab et al. 2004).
The early to late Paleozoic active continental
margin along the Kunlun range apparently initiated
a tectonic framework that influenced all subsequent paleogeographic developments in northern
Tibet (Xiao et al. 2002a; Roger et al. 2003). Heubeck
(2001) showed that this active margin extended
from the Qinling to the Caucasus from at least the
Middle Devonian to the Late Permian (fig. 8).
Therefore, in the late Paleozoic to early Mesozoic,
a continuous active margin of southern Eurasia extended for some 10,000 km north of the paleoTethys ocean (figs. 1, 8). This subduction-accretion
history was mostly superimposed on Paleozoic accretion and amalgamation along a northwarddipping subduction zone beneath the southern active margin of Eurasia (Lin et al. 1985; Reischmann
et al. 1990; Kröner et al. 1993; Meng and Zhang
1999, 2000; Heubeck 2001; Ratschbacher et al.
2003). The amalgamation of the southerly derived
blocks that accreted to Eurasia in MesozoicCenozoic time (including the Qiangtang-Cimmeria
block in the Late Jurassic) took place in this paleo-
geographic framework (Dewey et al. 1988; Kapp et
al. 2000, 2003a, 2003b).
ACKNOWLEDGMENTS
We are grateful to the personnel of the Beijing
SHRIMP Laboratory for their kind assistance. We
sincerely thank Q. L. Hou, Z. H. Wang, J. Hao, A.
M. Fang, G. C. Zhang, H. L. Chen, and H. Zhou for
their help in the field and laboratory. W. J. Xiao is
grateful to the University of Hong Kong, where he
was invited as a visiting scholar and prepared this
manuscript. Discussions with Y. S. Pan, W. M.
Deng, J. Aitchison, G. C. Zhao, and M.-F. Zhou
greatly improved early drafts. We thank two anonymous reviewers and A. Anderson for constructive
comments and suggestions that greatly improved
the manuscript. Funding by Projects of the Chinese
National Science Foundation (40172080, 40234045,
40032010, and 40372042), the Chinese Academy of
Sciences (KZCX2-SW-119), the Hong Kong Research Grants Council (HKU7040/04P), and the
Royal Society of London is gratefully acknowledged. This article forms a contribution to IGCP
473 and 420.
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