Precambrian Research 257 (2015) 47–64
Contents lists available at ScienceDirect
Precambrian Research
journal homepage: www.elsevier.com/locate/precamres
The 600–580 Ma continental rift basalts in North Qilian Shan,
northwest China: Links between the Qilian-Qaidam block and SE
Australia, and the reconstruction of East Gondwana
Xin Xu a , Shuguang Song a,b,∗ , Li Su c , Zhengxiang Li d , Yaoling Niu b , Mark B. Allen b
a
MOE Key Laboratory of Orogenic Belt and Crustal Evolution, School of Earth and Space Sciences, Peking University, Beijing 100871, China
Department of Earth Sciences, Durham University, Durham DH1 3LE, UK
c
Geological Lab Center, China University of Geosciences, Beijing 100083, China
d
ARC Center of Excellence for Core to Crust Fluid Systems (CCFS) and The Institute for Geoscience Research (TIGeR), Department of Applied Geology, Curtin
University, Perth 6845, Australia
b
a r t i c l e
i n f o
Article history:
Received 5 September 2014
Received in revised form
12 November 2014
Accepted 25 November 2014
Available online 4 December 2014
Keywords:
Zhulongguan basalts
Continental rifting
Late Neoproterozoic
Qilian-Qaidam block
Breakup of East Gondwana
a b s t r a c t
We report a sequence of thick, well-preserved basaltic lavas interlayered with shallow marine dolomitic
carbonates, mudstones and siltstones of the Zhulongguan Group, in the western segment of the North Qilian orogen, northwest China. Two new zircon SIMS ages show that this sequence formed at ∼600–580 Ma.
The mafic volcanics can be subdivided into tholeiitic and alkaline basalts, and have compositions similar to present-day ocean island basalt (OIB) or continental flood basalts. The occurrence, geochemical
features and age data suggest that the Zhulongguan basalts originated at a continental rift setting in the
latest Neoproterozoic, within the north margin of the Qilian-Qaidam block. This volcanic-sedimentary
formation exhibits close affinity to the passive continental margin in southeastern Australia. Our observations favor a link of the Qilian-Qaidam block with SE Australia (also south China) during the breakup
of Rodinia, thereby filling a void in existing reconstructions of the region.
© 2014 Elsevier B.V. All rights reserved.
1. Introduction
Intraplate magmatism, especially continental flood basalts
induced by mantle plumes or superplumes, plays an important
role in reconstructing the framework of supercontinents (White
and McKenzie, 1989; Hill et al., 1992; Saunders et al., 1996; Li
et al., 1999, 2008b; Ernst et al., 2008). There is a complete spectrum of within-plate magmatism from extensive sub-alkaline flood
basalt provinces to rift volcanism with more alkaline provinces
(Wilson, 1989). Syn-rift sedimentation often proceeds into continental breakup, when rifting ceased (i.e. the drift stage) and a
new ocean spreading center was created (e.g. Powell et al., 1994).
Therefore, comparison of geochemical fingerprints of key magmatic
events, together with lithostratigraphic correlation of contemporary rift successions, may help to establish the configuration of
ancient continental masses (Li et al., 2008b; Ernst et al., 2008).
∗ Corresponding author at: MOE Key Laboratory of Orogenic Belt and Crustal Evolution, School of Earth and Space Sciences, Peking University, Beijing 100871, China.
Tel.: +86 10 62767729.
E-mail address: [email protected] (S. Song).
http://dx.doi.org/10.1016/j.precamres.2014.11.017
0301-9268/© 2014 Elsevier B.V. All rights reserved.
The transition of the tectonic regime from the assembly of the
Neoproterozoic supercontinent Rodinia to its breakup is thought to
have occurred in the period of 0.9–0.86 Ga, which corresponds to
a magmatic quiescence in South China (Li et al., 2003, 2010a,b,c).
Multiple episodes of anorogenic magmatism during 850–720 Ma
are widely distributed in South China, Tarim, North America, India,
South Korea, Southern Africa, and Australia (Powell et al., 1994;
Park et al., 1995; Wingate et al., 1998; Preiss, 2000; Frimmel et al.,
2001; Lee et al., 2003; Li et al., 1999, 2003, 2008a,b, 2010a,b,c; Ling
et al., 2003; Wang and Li, 2003; Xu et al., 2005; Lu et al., 2008; Ernst
et al., 2008). They are believed to be associated with the breakup of
Rodinia, induced by mantle plumes or a superplume (Li et al., 1999,
2003, 2008b; Wang et al., 2007, 2008, 2009, 2010; Ernst et al., 2008).
In addition, there are geological records suggesting the separation
of microcontinents from the eastern Australia-east Antarctica continental margin during the 600–550 Ma interval (Crawford, 1992;
Veevers et al., 1997; Crawford et al., 1997; Wingate et al., 1998;
Foden et al., 2001; Direen and Crawford, 2003; Meffre et al., 2004;
Fergusson et al., 2009).
In the Qilian-Qaidam block between South China and Tarim,
within-plate magmatic rocks of 850–750 Ma have also been
recognized, including mafic-ultramafic intrusions, mafic dykes,
continental flood basalts, and anorogenic granites, and they were
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X. Xu et al. / Precambrian Research 257 (2015) 47–64
interpreted to be correlated with the fragmentation of Rodinia (Li
et al., 2005; Tseng et al., 2006; Lu et al., 2008; Song et al., 2010;
Tung et al., 2013).
The Early Paleozoic North Qilian Orogen (NQO) is located at
the northeastern margin of the Tibetan Plateau, NW China, within
the tectonically active Qilian Shan. It formed by the closure of a
Neoproterozoic to Early Paleozoic ocean and recorded a complete
Wilson Cycle from the continental breakup to collision of the
Qilian and Alxa basement blocks (for details, see Song et al., 2013
and references therein). The Qilian block is itself separated from
the larger Qaidam block by the Early Paleozoic North Qaidam UHP
belt, and they are suggested to be of Yangtze affinity on the basis of
Meso- to Neoproterozoic intrusions relevant to the amalgamation
of Rodinia supercontinent (Guo et al., 1999; Wan et al., 2001,
2006; Song et al., 2012; Tung et al., 2007, 2013). However, the
geological evolution of the Qilian-Qaidam block, especially its
relation with South China, Tarim and position in Rodinia during
the late Neoproterozoic, are still not constrained. The sparse
outcrop of within-plate magmatism hinders a direct comparison
with other fragments of Rodinia. This may be attributed to the
superimposed tectonic modification, the subduction of passive
continental margins or their deep burial during the later stages of
the orogeny (Yin et al., 2008; Song et al., 2014).
In this paper, we present new field observations, SIMS U–Pb
zircon ages, elemental and Sr–Nd isotopic data, and mineral
compositions for the basalts interbedded with shallow marine
sedimentary rocks in the northern margin of the Qilian-Qaidam
block. A better understanding of this volcanic sequence will
enable a useful comparison with the volcanic passive margin
in southeastern Australia and the Late Precambrian formation
in South China. Such studies will not only provide insights
into the development history of the North Qilian Orogen during the Precambrian, but may also reveal the relationship of the
Qilian-Qaidam block, South China and Australia in the context of
Rodinia.
2. Geological setting
The Qilian-Qaidam block in the northern Tibetan Plateau is
presently surrounded by three Precambrian cratons, i.e. the North
China Craton (NCC) to the east, the Tarim Craton (TC) to the northwest and the South China Craton (SCC) to the southeast (Fig. 1a).
It consists of the North Qilian oceanic suture zone, the Qilian
block, the North Qaidam UHP belt and the Qaidam block, from
north to south. The North Qilian oceanic suture zone (namely
North Qilian Orogen) extends NW–SE for ∼1000 km (Fig. 1b). In
the northwest, it is offset by a sinistral strike-slip Altyn Tagh
Fault (ATF) for up to 400 km and in direct contact with the Dunhuang block (Zhang et al., 2001). In the northeast, the Alxa (also
known as Alashan) block is bounded by the Longshoushan Fault
(LF), and considered to be the westernmost component of NCC
due to the similar Archean-Paleoproterozoic gneisses in the two
regions (Zhao and Cawood, 2012; Zhang et al., 2013a,b). However, the ∼827 Ma Jinchuan Cu–Ni-bearing ultramafic rocks and
800–900 Ma granitoids implied that the Alxa block may be a fragment of Rodinia, with affinities to the Qilian and South China
blocks in the Upper Proterozoic (Li et al., 2005; Song et al., 2013).
The Qilian block in the south is bounded on its northeast side by
the North Margin Fault (NMF) and has a Precambrian basement
which has the affinity with the Yangtze block, i.e. the northern
part of the larger South China block (Wan et al., 2001, 2006; Lu
et al., 2008; Song et al., 2010, 2012, 2013; Tung et al., 2007, 2013).
A Paleoproterozoic terrane, namely the Quanji Massif, was recognized in the south part of the Qilian block, which consists of
Paleoproterozoic granitic gneisses and mafic granulite with ages
of 2470–1800 Ma (Zhang et al., 2001; Chen et al., 2007, 2009a; Lu
et al., 2008).
Further south is the North Qaidam UHP metamorphic belt, representing a continent–continental collision zone along the northern
margin of the Qaidam Basin (e.g. Song et al., 2004a, 2014 and
references therein). The UHP belt is mainly consisted of granitic
gneisses, pelitic gneisses, with eclogites and garnet peridotites.
It is believed that the continental crust including orthogneisses
(1000–900 Ma) had subducted into depth of 200 km and exhumed
with enclosed UHPM rocks in the period of 460–400 Ma (Song
et al., 2012). Two episodes of orogeny during the Grenville and
Caledonian age, involved progression from oceanic subduction to
continental collision, have been confirmed by Song et al. (2013,
2014). The Qaidam Basin to the south is covered by a Mesozoic
to Cenozoic sediments and underlain mainly by Precambrian crystalline basement (Wan et al., 2006). The basement rocks mainly
exposed in the North Qaidam UHPM belt and south margin of the
Qaidam block (Song et al., 2014).
Previous works have reached a consensus that the Qilian and
Qaidam blocks have close affinities with South China according to
the orogenic and rifting events related to the assembly and breakup
of Rodinia, respectively (Guo et al., 1999; Lu et al., 2008; Song et al.,
2012, 2013; Tung et al., 2007, 2013). Further Wan et al. (2001,
2006) emphasized that the high-grade basement of the North Qilian orogenic belt has similar Nd isotopic compositions with those
of the North Qaidam UHPM belt. Thus the Qilian and Qaidam blocks
form one integrated terrane, i.e. the Qilian-Qaidam block during the
Precambrian.
The North Qilian Orogen (NQO) is one of the best preserved oceanic-type cold subduction belts in China, resulting
from closing of the ancient Qilian Ocean between Alxa and the
Qilian-Qaidam block during the Early Paleozoic (Xiao et al., 1978;
Wu et al., 1993; Feng and He, 1996; Zhang et al., 2007; Song
et al., 2004b, 2006, 2007, 2009, 2013; Xiao et al., 2009; Chen
et al., 2014). It consists dominantly of Middle-Late Proterozoic
high-grade metamorphic basement, Late Proterozoic low-grade
metamorphic volcanic and sedimentary successions, Early Paleozoic subduction-related rock associations (ophiolite complexes,
high-pressure/low-temperature metamorphic rocks, arc-related
volcanics and intrusions), Silurian flysch and Devonian molasse formations, and later sedimentary cover (Fig. 1b). The present-day
high topography of the NQO results from the India-Asia collision and Tibetan plateau uplift in the late Cenozoic (Yin et al.,
2008).
Precambrian fragments in the northern margin of the Qilian
block have been juxtaposed with arc rocks during the Early Paleozoic collision-accretion process (Fig. 1b). In the western segment,
the stratigraphic succession contains a pre-Sinian group, Sinian
volcanic-sedimentary sequence and Early Paleozoic cover strata
(Fig. 2b).
The 900–1000 Ma orthogneisses constitute the oldest and major
component of the Precambrian basement of the Qilian-Qaidam
block (Guo et al., 1999; Wan et al., 2001; Li et al., 2007; Song et al.,
2012; Tung et al., 2007, 2013). In addition, within-plate magmatism
(850–750 Ma) including diabasic dyke swarms, mafic-ultramafic
intrusions, anorogenic granitoids, and the remnants of continental flood basalts, have been recognized (Li et al., 2005; Tseng et al.,
2006; Lu et al., 2008; Song et al., 2010; Tung et al., 2013).
The Zhulongguan Group mainly crops out in the northwestern
part of the Qilian block with total area of more than 1000 km2 and a
thickness of about 3–7 km (Fig. 2a). The volcano-sedimentary succession belt is controlled by regional-scale faults and extends along
the main axis of the NQO. This group consists predominantly of
the low-grade metamorphic volcanic layers interbedded with shallow marine dolomitic limestone, terrigenous and pyroclastic rocks
and iron-bearing quartzite (Xia et al., 2000), which constitute a
X. Xu et al. / Precambrian Research 257 (2015) 47–64
49
Fig. 1. (a) Tectonic location of the Qilian-Qaidam blocks in NW China (after Song et al., 2013); (b) simplified geological map of the Qilian-Qaidam region. Abbreviations: ATF
– Altyn Tagh Fault, LF – Longshoushan Fault, NMF – North Margin Fault of the Qilian block.
multi-cycle volcanic-sedimentary succession (Fig. 2b). This group is
tectonically juxtaposed with the Aoyougou ophiolite (495–504 Ma,
Xiang et al., 2007; Song et al., 2013; Fig. 2c) and intruded by a 430 Ma
adakite pluton (Chen et al., 2012).
The overlying sequences, namely the Jingtieshan and Daliugou groups, are faulted against the Zhulongguan Group and
are dominantly constituted of sandstones, siltstones, mudstones,
dolomites interbedded with mafic volcanic and iron ore layers
(Fig. 2b), which were considered as the middle and upper part of
the Zhulongguan Group (Xia et al., 2000). The Baiyanggou Group
in the uppermost Sinian succession is recognized as a suite of
thick coarse clastic rocks, including tillitic and sandy conglomerates. The pebbles from the basal tillite are derived from the
underlying Jingtieshan and Daliugou Group, which indicates rapid
accumulation during the rifting stage (Zuo et al., 1999). The Early
Paleozoic complex is dominantly consisted of Cambrian to Ordovician arc-related volcanic and sedimentary rocks and Silurian flysch
formation.
The regional importance of these Late Precambrian rocks is
sometimes downplayed, such that the entire Qilian Shan is referred
to as an accretionary orogenic belt without significant Precambrian
crust (e.g. Şengör, 1990; Xiao et al., 2009). However, the extent,
thickness, continuity and stratigraphy of the Precambrian succession, the absence of major metamorphism and presence of ∼1 Ga
continental basement, all point to a microcontinental terrane(s) of
sufficient size to be considered in regional and global plate reconstructions.
3. Petrography of the Zhulongguan basalts
Samples were collected from three sections in the Zhulongguan Group (see localities in Fig. 2a). Two representative
sections, rock assemblages and field relations are shown in
Fig. 2c.
In the Aoyougou valley (Section 1), four layers of mafic volcanic lavas are interbedded with Precambrian carbonate layers;
they constitute multiple eruption–deposition cycles. The basaltic
lavas can reach up to 300 m in thickness. They are weakly altered
and have massive (locally pillow), vesicular/amygdaloidal structures (Fig. 3b and d). Some of them are porphyritic with abundant
plagioclase and augite phenocrysts in a usually intersertaltextured groundmass filled with plagioclase laths, chloritised
glass and Fe–Ti oxides. Most of these lavas show ophitic texture
with euhedral plagioclase skeletons and subhedral augite grains
(Fig. 3g).
Massive basalt samples (11QL-65 and 66) come from the lower
part of the Zhulongguan Group near Qiqing village (Section 2
in Fig. 2). The volcanic interlayers are 100–300 m in thickness,
interbedded with volcanic breccias, tuffs, siliceous slates, siltstone/sandstone, and limestone. In thin sections, they mainly
consist of clinopyroxene, plagioclase and minor alteration minerals
(actinolite–chlorite).
The rest of the samples were collected from the Jiugeqingyang
section (Fig. 2). These basaltic lavas are accompanied by iron ore
beds, sandstone, tuff, volcanic breccia, shale and pelite. The volcaniclastic sample (13QL-18) consists of detrital components including
rock fragments (basaltic glass) and mineral clasts (clinopyroxene
and olivine) (Fig. 3h). Some of basaltic samples show ophitic texture
similar to some of the Aoyougou section while the others contain abundant altered clinopyroxene and plagioclase phenocrysts
with the groundmass glass replaced by secondary chlorite, epidote and calcite (Fig. 3e and f). The clinopyroxene phenocrysts
are commonly subhedral to anhedral with a diopside composition
(Wo46–48 En38–42 Fs12–16 ).
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X. Xu et al. / Precambrian Research 257 (2015) 47–64
Fig. 2. (a) Schematic geological map of the Qiqing area (modified after Xia et al., 2012). (b) Stratigraphic column of the Qiqing area, showing the Precambrian basement,
Neoproterozoic rifting-related volcanic-sedimentary succession and Paleozoic orogenic complex. (c) Two cross-sections of the Zhulongguan Group with sample localities.
4. Analytical methods
Zircons were separated from 11QL-65 and 13QL-18 by using
standard density and magnetic separation techniques. Zircon
grains, together with the standard zircon Plésovice and Qinghu,
were embedded in an epoxy mount and then polished down to
expose the inner structure for analysis. The CL examination was
done by using a FEI QUANTA650 FEG Scanning Electron Microscope
(SEM) under conditions of 15 kV/120 nA in the School of Earth and
Space Science, Peking University, Beijing.
Measurements of U–Th–Pb isotopes were conducted using a
Cameca IMS-1280 SIMS in the Institute of Geology and Geophysics,
Chinese Academy of Sciences in Beijing. The instrument description
and analytical procedure is given in Li et al. (2009). The primary O2 −
ion beam spot is about 20–30 mm in size. Analysis of the standard
zircon Plésovice was interspersed with analysis of unknowns. Each
measurement consists of 7 cycles. Pb/U calibration was performed
relative to zircon standard Plésovice (337 Ma, Sláma et al., 2008);
U and Th concentrations were calibrated against zircon standard
91,500 (Wiedenbeck et al., 1995). A long-term uncertainty of 1.5%
(1 RSD) for 206 Pb/238 U measurements of the standard zircons was
propagated to the unknowns (Li et al., 2010a,b,c), despite that the
measured 206 Pb/238 U error in a specific session is generally 1% (1
RSD). Measured compositions were corrected for common Pb using
X. Xu et al. / Precambrian Research 257 (2015) 47–64
51
Fig. 3. Field and photomicrographs of the Zhulongguan basalts. (a) Tholeiitic basalt conformably contacting dolomitic limestones. (b) The massive basalt with amygdalae. (c)
The thick basaltic lava layers with siltstone. (d) The thick pillow lavas. (e, f) Clinopyroxene phenocrysts in alkaline basalts (12QL-101). (g) The intersertal and ophitic texture
showing pyroxene grains within the plagioclase skeletons in tholeiitic basalt (12QL-106). (h) Basaltic glass and detrital minerals (olivine and clinopyroxene) in volcaniclastic
sample (13QL-18).
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X. Xu et al. / Precambrian Research 257 (2015) 47–64
non-radiogenic 204 Pb. Corrections are sufficiently small to be insensitive to the choice of common Pb composition, and an average
of present-day crustal composition (Stacey and Kramers, 1975) is
used for the common Pb assuming that the common Pb is largely
surface contamination introduced during sample preparation. Data
reduction was carried out using the Isoplot/Ex v. 2.49 program
(Ludwig, 2001). Uncertainties on individual analyses in data tables
are reported at 1 level; Concordia U–Pb ages are quoted with 95%
confidence interval.
In order to monitor the external uncertainties of SIMS U–Pb zircon dating calibrated against Plésovice standard, an in-house zircon
standard Qinghu was alternately analyzed as an unknown together
with other unknown zircons. The measurements on Qinghu zircon yield Concordia ages of 160.2 ± 0.8 Ma and 159.3 ± 1.9 Ma,
which are identical within error with the recommended value of
159.5 ± 0.2 Ma (Li et al., 2013a).
Bulk-rock major element oxides (SiO2 , TiO2 , Al2 O3 , FeO, MnO,
MgO, CaO, Na2 O, K2 O, and P2 O3 ) were determined using inductively
coupled plasma-atomic emission spectroscopy (ICP-OES) at China
University of Geosciences, Beijing. The analytical uncertainties are
generally less than 1% for most elements with the exception of TiO2
(∼1.5%) and P2 O5 (∼2.0%). The loss on ignition was measured by
placing 1 g of powder in the furnace at 1000 ◦ C for several hours
before cooled in a desiccator and reweighted. The trace element
analysis for Zhulongguan basalt samples were accomplished on
an Agilent-7500a inductively coupled plasma mass spectrometer
(ICP-MS) at China University of Geosciences, Beijing. The detailed
analytical procedures follow Song et al. (2010). The relative difference between measured and recommended values for two USGS
rock reference materials (BCR-1 and BHVO-1) indicates that analytical accuracy is better than 5% for most elements, ranging between
10% and 13% for Cu, Sc, Nb, Er, Th, and U, and between 10% and 15%
for Ta, Tm, and Gd.
The bulk-rock Sr–Nd isotope analyses are accomplished at
MOE Key Laboratory of Orogenic Belts and Crustal Evolution,
Peking University. About 300 mg unknown samples and ∼200 mg
standard samples (BCR-2) were dissolved by using HF + HNO3
in Teflon vessels and heated at 140 ◦ C for a week in order to
be completely dissolved. The pure Sr and Nd were obtained
by passing through conventional cation columns (AG50W and
P507) for analysis using a multi-collector inductively coupled
plasma mass spectrometer (MC-ICP-MS) of the type VG AXIOM.
Mass fractionation corrections for Sr and Nd isotopic ratios were
normalized to 86 Sr/88 Sr = 0.1194 and 146 Nd/144 Nd = 0.7219, respectively. Repeated analyses for the Nd and Sr standard samples
(JNdi and NBS987) yielded 143 Nd/144 Nd = 0.512120 ± 11 (2) and
87 Sr/86 Sr = 0.710250 ± 11 (2), respectively.
5. Results
5.1. U–Pb zircon age
One basalt (11QL-65) and one volcaniclastic sample (13QL-18)
were selected for SIMS zircon U–Pb dating. Sample 11QL-65 comes
from the lower half of the succession, and sample 13QL-18 from
near the top (Fig. 2). The results are listed in Table 1 and illustrated on concordia plots in Fig. 4B and D. Zircon grains from
11QL-65 show mostly irregular and fragmentation tabular shapes
with length up to 100 ␮m and length-width ratios up to 2. The CL
images (Fig. 4A) display slight to dark luminescence and homogeneous structure with straight and wide growth bands, which are
similar to zircons from mafic volcanic and gabbroic rocks (Song
et al., 2010). The CL images of zircons from 13QL-18 are similar to
those from 11QL-65. As shown in Fig. 4C, these grains are mainly
irregular crystals, indicating that they are directly derived from the
Zhulongguan basalts.
The zircons of the basalt sample (11QL-65) have various
abundances of Th (43–2158 ppm) and U (90–1340 ppm) with
relatively high Th/U ratios (0.47–1.6). Eleven analyses yield apparent 206 Pb/238 U ages of 571–626 Ma and form a concordia age
of 600 ± 7 Ma (MSWD = 0.14) (Fig. 4B); the other three spots
give distinctly older 207 Pb/206 Pb apparent ages of 1471 ± 25 Ma,
1994 ± 13 Ma and 2520 ± 9 Ma, respectively. The uniform CL images
and U–Pb ages suggest that the first group of zircons was crystallized from a basaltic magma and could represent the eruption time
of Zhulongguan basaltic lava, whereas the old zircons (1.5–2.5 Ga)
may be derived from the Precambrian basement. Similarly, the uranium content in zircons from sample 13QL-18 varies in a large range
from 92 to 818 ppm and Th from 50 to 1041 ppm with Th/U ratios
of 0.42–1.27. One spot (#10) was excluded for its high common
Pb (f206 = 1.16), and other nine analyses yield 206 Pb/238 U apparent
ages ranging from 567 to 597 Ma and a weighted average age of
583 ± 7 Ma (MSWD = 1.4), the same as the concordia U–Pb age of
583 ± 3 Ma (n = 9, MSWD = 0.72). In conclusion, the Zhulongguang
basalts were formed at ∼600–580 Ma.
5.2. Geochemistry
5.2.1. Whole-rock major and trace elements
Fifteen basalt samples (see localities in Fig. 2) were analyzed for
major and trace elemental compositions (Table 2). All the analyses are plotted on an anhydrous basis (Fig. 5). Of these samples,
5 plot in the alkaline field and 10 in the subalkaline basalt field
on the Nb/Y versus Zr/TiO2 diagram (Fig. 5a). All the subalkaline
basalts belong to the tholeiitic series in the FeOt/MgO versus TiO2
plot (Fig. 5b). The tholeiitic samples have low Ti/Y ratios (<500),
whereas alkaline samples show high Ti/Y ratios (>500) based on
the classification of Xu et al. (2001). All of the mafic rocks are sodic
series (Na2 O > K2 O).
The tholeiitic basalts have relatively high SiO2 (48–56%), Fe2 O3 T
(13.5–17.8%), Y (34.8–46.8 ppm) and HREE, but low MgO (3.1–6.0%),
Mg# (31–45) and compatible elements. The low contents of MgO, Cr
(12–83 ppm) and Ni (20–64 ppm) are far from the expected composition of melts in equilibrium with the mantle peridotite (Cox,
1980; Wilson, 1989), indicating significant fractional crystallization
and/or crustal contamination.
The alkaline basalts show a relative narrow compositional
variation with lower SiO2 (49–52%), TiO2 (1.5–3.2%) and Fe2 O3 T
(11.7–15.5%). The higher MgO (6.7–11.1%), Mg# (53–69), Cr (up to
622 ppm) and Ni (up to 280 ppm) imply less evolved features. Sample 11QL-65 has the highest MgO (11.1%), Mg# (69), Cr (622 ppm)
and Ni (280 ppm) which is similar to the picritic or primitive highMg (e.g. Mg# > 65 and/or MgO > 9 wt.%) magma.
On primitive-mantle normalized multiple trace elements diagrams, the alkaline basalts have the uniform “humped” distribution
patterns characterized by variable enrichment in Rb, Ba, Pb, Nb, Ta,
Nd and Ti, and depletion in Th, U, Sr, P and Y, which are akin to
those of the high-Ti picritic basalts in Deccan Traps of India (Melluso
et al., 2006; Fig. 6c). On the contrary, the tholeiitic basalts display
uniform negative anomalies of HFSE (Nb* = 0.3–0.9; Niu and Batiza,
1997), various depletion of Sr, P, Eu, and positive anomalies of most
incompatible elements including Th, U and Ti with a large variation
in Rb and Ba. The composition of the tholeiites is similar to the lowTi basalts in several large igneous provinces, such as Emeishan in
South China (Xu et al., 2001; Fig. 6d).
As shown in Fig. 6a and b, all the samples exhibit consistent LREE
enrichment ((La/Yb)n = 5.4–7.2, 2.3–4.5 for the alkaline and tholeiitic basalts, respectively). Most alkaline samples shows positive
anomalies of Eu (Eu/Eu* = 1.03–1.70, except for 12QL-109), whereas
the latter has the uniform Eu depletion (Eu/Eu* = 0.77–0.94). These
features are consistent with the observation that the tholeiitic
basalts are more evolved than the alkaline group. In summary, the
Table 1
SIMS zircon U–Pb data for the Zhulongguan basalt (11QL-65) and volcaniclastic rock (13QL-18).
Spot#
U ppm
Th ppm
Th/U
f206 (%)
207
206
Pb
Pb
±1 (%)
235
116
851
103
200
127
984
215
122
1763
2158
595
52
43
209
0.51
1.15
0.66
0.72
0.64
1.20
0.49
0.32
1.39
1.61
0.74
0.54
0.47
0.72
0.20
0.03
0.00
0.09
0.00
0.23
0.05
0.04
0.08
0.06
0.17
0.38
0.83
0.17
0.0922
0.0602
0.0653
0.0608
0.0588
0.0590
0.1226
0.1662
0.0600
0.0596
0.0584
0.0602
0.0603
0.0591
1.33
1.64
2.76
2.31
2.73
1.48
0.76
0.55
1.06
1.28
1.63
5.51
8.84
2.34
13QL-18
1
2
3
4
5
6
7
8
9
10
325
308
50
343
123
82
90
331
1041
124
0.85
0.81
0.55
0.83
0.66
0.52
0.42
1.11
1.27
0.55
0.00
0.05
0.00
0.36
0.45
0.13
0.02
0.00
0.01
1.16
0.05868
0.06004
0.05988
0.05964
0.06056
0.05872
0.05850
0.05924
0.05930
0.05616
0.73
0.91
1.49
0.99
1.66
1.17
1.35
0.95
0.53
2.77
Pb
U
3.26
0.82
0.87
0.78
0.79
0.78
5.99
10.66
0.82
0.84
0.80
0.80
0.82
0.80
0.78453
0.80147
0.78667
0.78620
0.76781
0.75873
0.77238
0.76865
0.76770
0.71487
±1? (%)
206
238
Pb
U
±1 (%)
207/206 Age (Ma)
±1
207/235 Age (Ma)
±1
206/238 Age (Ma)
±1
2.01
2.23
3.14
2.75
3.12
2.13
1.68
1.61
1.83
1.97
2.23
5.71
8.97
2.78
0.257
0.099
0.097
0.093
0.097
0.096
0.354
0.465
0.099
0.102
0.099
0.096
0.098
0.098
1.51
1.50
1.50
1.50
1.52
1.52
1.51
1.51
1.50
1.50
1.52
1.50
1.53
1.50
1471
611
784
632
559
566
1994
2520
605
590
546
611
614
570
25
35
57
49
58
32
13
9
23
28
35
115
180
50
1472
610
636
583
589
587
1974
2494
607
619
596
597
606
597
16
10
15
12
14
10
15
15
8
9
10
26
42
13
1473
610
595
571
597
592
1955
2463
607
626
609
594
604
604
20
9
9
8
9
9
25
31
9
9
9
9
9
9
1.67
1.75
2.11
1.85
2.24
1.90
2.02
1.79
1.60
3.24
0.0970
0.0968
0.0953
0.0956
0.0919
0.0937
0.0958
0.0941
0.0939
0.0923
1.50
1.50
1.50
1.56
1.50
1.50
1.50
1.52
1.51
1.68
555
605
599
591
624
557
549
576
578
459
16
20
32
21
35
25
29
21
12
60
588
598
589
589
579
573
581
579
578
548
8
8
10
8
10
8
9
8
7
14
597
596
587
589
567
577
590
580
579
569
9
9
8
9
8
8
9
8
8
9
X. Xu et al. / Precambrian Research 257 (2015) 47–64
11QL-65
227
1
740
2
158
3
278
4
198
5
817
6
440
7
379
8
1271
9
1340
10
801
11
97
12
90
13
290
14
380
381
92
413
188
159
217
297
818
224
207
f206 is the percentage of common 206 Pb in total 206 Pb. All error is 1sigma (1).
53
54
X. Xu et al. / Precambrian Research 257 (2015) 47–64
Fig. 4. (A and B) Cathodoluminescence images of representative zircons; (C and D) Concordia plot for sample 11QL-65 and 13QL-18.
Zhulongguan basalts show immobile trace elements characteristics similar to the present-day OIB and/or at least some continental
flood basalts, such as the Emeishan and Deccan lavas.
5.2.2. Whole-rock Sr–Nd isotopic data
Five alkaline and seven tholeiitic basalts were analyzed for
whole-rock Sr–Nd isotopic composition. The results are presented in Table 3 and illustrated in Fig. 7. The initial values
of the Sr–Nd isotope were calculated at 600 Ma. The alkaline
basalts have low 87 Sr/86 Sr ratios (0.70736–0.70848) and high
143 Nd/144 Nd ratios (0.512656–0.512733). In spite of the deviation from the mantle array due to the high initial Sr isotopic
values, the positive εNd values (4.1–5.3) are similar to those
of modern plume-related basalts and high-Ti basalts in several
famous LIPs (Fig. 7A). It is notable that they are also identical to that of picritic and upper basaltic volcanics on King
Island, Tasmania (εNd (579 Ma) = +3.5 to +4.8; Meffre et al., 2004)
and the high-Nb basalts of Mt Arrowsmith and Wright in New
South Wales (εNd (586Ma) = +3.7 to +4.7; Crawford et al., 1997)
(Fig. 7B).
On the contrary, the tholeiitic basalts have low 143 Nd/144 Nd values ranging from 0.512312 to 0.512695 and high 87 Sr/86 Sr values
from 0.70865 to 0.71977. The extremely high Sr isotopic values may
be attributed to the alteration of sea water. In general, the Sr–Nd
isotopic characteristics of the tholeiitic basalts are alike to those
of low-Ti basalts from Emeishan, Deccan and Siberia (Fig. 7A). All
tholeiitic samples show similar Sm–Nd isotopic compositions to
those of Eastern Australia volcanics (Meffre et al., 2004) (Fig. 7B).
6. Discussion
6.1. Petrogenesis
Primary melt composition not only reflects the pressure and
temperature conditions during partial melting, but also the compositions of source from which they derived (Putirka, 2005; Putirka
et al., 2007; Herzberg et al., 2007; Herzberg and Asimow, 2008; Niu
and O’Hara, 2008; Lee et al., 2009; Humphreys and Niu, 2009; Niu
et al., 2011; Wang et al., 2012). Nevertheless, magmas are the integrated products of the dynamic melting regime and complicated
melt transport process (Wilson, 1989; Niu and O’Hara, 2008). Thus
we need to evaluate the effect of later shallow level processes such
as fluid alteration and AFC (assimilation and fractional crystallization) process on the elemental abundance and isotopic ratios, prior
to an analysis of the potential mantle source. The elemental mobility can been estimated by the correlation between Zr (immobile in
the fluids alteration) and other elements (Wang et al., 2008). For the
Zhulongguan basalts, the high field strength elements (Nb, Ta, Ti,
Zr, Hf), REE, V, Th, U and Sr are essentially immobile during metamorphism and alteration. On the other hand, CaO, Na2 O, K2 O, Ba,
Rb and Pb show no linear relation with zirconium. Therefore these
mobile elements must be excluded to discuss rock classification
and petrogenesis.
6.1.1. Fractional crystallization
The Zhulongguan basalts show a large variation in MgO, Mg# and
compatible trace elements, suggesting that they have undergone
X. Xu et al. / Precambrian Research 257 (2015) 47–64
55
Table 2
Whole-rock major and trace element data for the Zhulongguan basalts.
Sample
Alkaline basalts
11QL-65
Major elements (wt.%)
50.01
SiO2
1.49
TiO2
10.86
Al2 O3
Fe2 O3 T
11.35
MnO
0.17
10.70
MgO
9.20
CaO
1.94
Na2 O
0.45
K2 O
P2 O5
0.18
LOI
3.14
Mg#
68.7
99.48
Total
Trace elements (ppm)
Sc
31.3
V
250
Cr
623
48
Co
280
Ni
Rb
14.32
221
Sr
17.2
Y
107
Zr
31.7
Nb
Ba
279
11.8
La
Ce
24.9
3.25
Pr
14.2
Nd
Sm
3.41
1.18
Eu
3.58
Gd
0.549
Tb
Dy
3.33
Ho
0.644
1.80
Er
Tm
0.243
Yb
1.55
Lu
0.226
Hf
2.70
Ta
1.21
Pb
1.31
Th
1.26
U
0.279
651
Ti/Y
Sample
Tholeiitic basalts
11QL-66
49.29
1.92
13.90
12.84
0.20
6.99
7.12
2.30
2.27
0.28
2.28
55.9
99.39
28.7
235
140
37
69
19.39
314
21.8
131
32.9
9474
15.7
32.7
4.19
18.1
4.27
2.59
5.07
0.717
4.28
0.827
2.24
0.302
1.92
0.273
3.10
1.50
1.14
1.45
0.368
622
12QL-101
12QL-107
46.82
2.26
14.45
13.29
0.17
6.41
7.41
3.10
1.17
0.33
4.73
52.9
100.13
47.47
1.92
13.45
12.00
0.30
7.36
4.23
2.70
1.70
0.26
8.62
58.8
100.00
33.3
413
65
49
71
19.20
252
21.4
119
33.3
1864
18.7
41.9
5.45
23.4
5.43
2.26
5.51
0.794
4.62
0.883
2.32
0.302
1.87
0.267
3.04
2.19
1.66
1.73
0.428
663
33.6
300
171
44
75
20.70
290
24.3
117
28.8
1654
17.3
37.7
4.65
19.2
4.54
1.77
5.03
0.770
4.76
0.964
2.64
0.355
2.26
0.325
2.86
1.59
4.11
2.08
0.499
506
12QL-109
09AY-01
09AY-07
09AY-09
09AY-10
09AY-11
46.59
2.96
12.98
14.30
0.20
7.46
3.67
3.65
0.13
0.33
7.69
54.9
99.98
46.90
2.80
12.50
14.60
0.17
2.80
8.10
1.58
1.14
0.22
9.10
30.9
99.91
51.22
2.76
11.83
16.09
0.15
5.70
5.32
2.85
0.82
0.21
2.95
45.2
99.89
54.08
3.09
13.15
12.98
0.27
4.54
4.39
1.99
1.68
0.25
3.46
44.9
99.88
52.05
2.68
12.18
17.12
0.16
4.54
6.28
1.58
0.87
0.20
2.23
38.2
99.90
54.14
2.88
11.76
15.75
0.16
3.73
6.83
2.04
0.71
0.23
1.68
35.5
99.89
31.7
463
31
52
44
2.89
276
29.7
158
42.9
228
20.3
46.0
6.00
26.0
6.22
2.04
6.58
0.994
5.97
1.182
3.18
0.427
2.68
0.383
4.07
2.74
1.69
2.11
0.507
669
41.7
477
61
43
54
41.74
187
37.2
155
13.4
107
14.7
33.8
4.46
19.0
4.84
1.33
5.71
0.930
6.17
1.340
3.91
0.570
3.78
0.550
3.62
0.76
5.87
3.21
0.850
408
40.3
496
12
40
21
14.33
149
40.1
164
12.7
157
20.8
45.1
5.66
23.2
5.64
1.58
6.34
1.025
6.54
1.335
4.01
0.574
3.78
0.565
3.80
0.72
6.63
4.93
1.709
384
48.2
563
14
72
26
45.52
152
42.2
203
14.3
289
20.7
47.6
6.12
24.8
5.93
1.95
6.80
1.065
6.63
1.347
3.84
0.520
3.27
0.471
4.21
0.78
7.53
5.63
1.494
413
40.0
479
13
37
20
28.10
163
34.9
154
11.9
156
17.7
40.2
5.07
20.8
5.06
1.45
5.76
0.911
5.77
1.183
3.45
0.501
3.33
0.481
3.61
0.68
4.98
4.61
1.164
409
40.7
467
12
41
19
10.57
162
42.7
183
13.8
223
22.3
48.8
6.17
25.2
5.94
1.68
6.88
1.099
6.90
1.437
4.21
0.594
3.96
0.580
4.28
0.81
7.35
5.56
1.414
368
Tholeiitic basalts
09AY-11
09AY-12
09AY-13
54.14
2.88
11.76
15.75
0.16
3.73
6.83
2.04
0.71
0.23
1.68
35.5
99.89
51.59
2.14
12.73
14.03
0.26
4.99
7.84
2.37
1.71
0.18
2.04
45.3
99.88
Trace elements (ppm)
Sc
40.7
467
V
12
Cr
Co
41
19
Ni
10.57
Rb
162
Sr
42.7
Y
41.8
383
59
43
50
42.44
212
34.8
Major elements (wt.%)
SiO2
TiO2
Al2 O3
Fe2 O3 T
MnO
MgO
CaO
Na2 O
K2 O
P2 O5
LOI
Mg#
Total
12QL-104
12QL-105
12QL-106
51.50
2.65
12.02
16.05
0.23
4.65
8.99
1.40
0.93
0.23
1.24
40.3
99.90
50.58
2.63
11.72
14.99
0.17
4.85
5.67
2.49
0.53
0.29
6.07
43.0
100.00
45.99
3.05
11.26
16.92
0.26
5.63
8.24
2.91
0.28
0.32
5.28
43.7
100.16
51.82
2.70
11.53
16.13
0.17
4.75
7.98
1.89
0.58
0.29
2.26
40.7
100.09
40.2
459
52
44
42
18.06
196
41.2
47.6
525
72
47
60
33.58
166
42.4
46.1
528
80
50
60
12.03
121
46.8
45.6
492
83
49
64
18.75
134
40.1
56
X. Xu et al. / Precambrian Research 257 (2015) 47–64
Table 2
Whole-rock major and trace element data for the Zhulongguan basalts.
Sample
Tholeiitic basalts
Zr
Nb
Ba
La
Ce
Pr
Nd
Sm
Eu
Gd
Tb
Dy
Ho
Er
Tm
Yb
Lu
Hf
Ta
Pb
Th
U
Ti/Y
09AY-12
09AY-13
12QL-104
12QL-105
12QL-106
183
13.8
223
22.3
48.8
6.17
25.2
5.94
1.68
6.88
1.099
6.90
1.437
4.21
0.594
3.96
0.580
4.28
0.81
7.35
5.56
1.414
368
146
11.2
311
14.2
32.5
4.23
17.8
4.46
1.28
5.24
0.847
5.38
1.105
3.23
0.469
3.16
0.461
3.14
0.62
4.00
3.39
0.879
360
166
14.2
174
16.1
37.1
4.81
20.8
5.28
1.58
6.27
1.028
6.69
1.393
4.08
0.582
3.91
0.574
3.74
0.79
4.32
3.43
0.907
368
207
19.9
73
14.1
37.6
5.40
24.8
6.93
2.15
8.05
1.299
8.32
1.718
4.85
0.677
4.45
0.645
5.38
1.36
3.50
2.81
0.727
409
212
21.3
118
16.0
40.0
5.65
26.0
7.27
2.20
8.55
1.393
8.87
1.859
5.24
0.729
4.73
0.688
5.53
1.53
3.17
2.77
0.685
399
194
19.1
104
13.8
35.7
5.06
23.3
6.54
1.96
7.58
1.232
7.91
1.637
4.64
0.650
4.26
0.622
5.10
1.27
3.06
2.50
0.662
423
4
Tholeiitic basalts
Alkaline basalts
a
b
Phonolite
Com/Pant
1
3
Rhyolite
Trachyte
0.1
TiO 2
Zr/TiO 2 *0.0001
10
09AY-11
Rhyodacite/Dacite
TrachyAnd
2
T
Andesite
0.01
Bsn/Nph
Andesite/Basalt
SubAlkaline
Basalt
0.001
0.01
1
Calc
le
ho
-Ala
iit
i
e
cs
kline
Alkaline Basalt
rie
s
serie
s
0
1
0.1
0
10
1
2
3
4
5
FeOt/MgO
Nb/Y
Fig. 5. (a) Nb/Y versus Zr/TiO2 × 0.0001 diagram (Winchester and Floyd, 1976). (b) FeOt/MgO versus TiO2 diagram (Miyashiro, 1974).
Table 3
Whole-rock Sr–Nd isotopic data for the Zhulongguan basalts.
Sr (ppm)
87
Alkaline basalts
14.32
11QL-65
19.39
11QL-66
19.20
12QL-101
20.70
12QL-107
2.89
12QL-109
221.0
313.8
251.8
290.2
275.6
0.1831
0.1745
0.2154
0.2015
0.0296
Tholeiitic basalts
41.74
09AY-01
45.52
09AY-09
28.10
09AY-10
09AY-12
42.44
33.58
12QL-104
12.03
12QL-105
18.75
12QL-106
187.0
151.7
163.2
211.6
166.5
121.4
134.4
0.6305
0.8478
0.4864
0.5665
0.5697
0.2798
0.3941
Rb (ppm)
Rb/86 Sr
87
Sr/86 Sr
2
ISr
Sm (ppm)
Nd (ppm)
147
Sm/144 Nd
0.707389
0.707357
0.708482
0.708023
0.707594
0.000010
0.000009
0.000016
0.000012
0.000019
0.70582
0.70586
0.70664
0.70630
0.70734
3.41
4.27
5.43
4.54
6.22
14.18
18.12
23.36
19.25
25.96
0.1527
0.1494
0.1476
0.1495
0.1520
0.712885
0.719766
0.715721
0.714250
0.712088
0.710405
0.708646
0.000019
0.000017
0.000224
0.000011
0.000019
0.000013
0.000018
0.70749
0.71251
0.71156
0.70940
0.70721
0.70801
0.70527
4.84
5.93
5.06
4.46
6.93
7.27
6.54
18.98
24.82
20.83
17.78
24.76
26.04
23.34
0.1618
0.1517
0.1540
0.1593
0.1776
0.1771
0.1778
Note: (1) ISr = 87 Sr/86 Sr − 87 Rb/86 Sr × (eT − 1), where Rb = 1.42 × 10−11 year−1 (Steiger and Jäger, 1977).
εNd
(T) = {[143 Nd/144 Nd − 147 Sm/144 Nd × (eT − 1)]/[(143 Nd/144 Nd)CHUR(0) − (147 Sm/144 Nd)CHUR(0) × (eT − 1)] − 1} × 10,000,
(2)
(143 Nd/144 Nd)CHUR(0) = 0.512638; (147 Sm/144 Nd)CHUR(0) = 0.1967 (Lugmair and Marti, 1978).
(3) T = 600 Ma, crystallization age of the Zhulongguan Group basalts.
2
εNd (T)
0.512696
0.512690
0.512656
0.512663
0.512733
0.000015
0.000009
0.000019
0.000019
0.000018
4.5
4.6
4.1
4.1
5.3
0.512491
0.512322
0.512312
0.512448
0.512680
0.512695
0.512686
0.000019
0.000018
0.000019
0.000018
0.000017
0.000015
0.000015
-0.2
-2.7
-3.1
-0.8
2.3
2.6
2.4
143
Nd/144 Nd
where
Sm = 6.54 × 10−12 year−1 ;
X. Xu et al. / Precambrian Research 257 (2015) 47–64
1000
57
1000
c
a
Rock/Chondrite
OIB
Rock Primitive mantle
Alkaline basalts
Deccan high-Ti pricite
100
10
1
Alkaline b asalts
Deccan h igh-Ti pricite
100
OIB
10
1
La
Ce
Pr
Nd
Sm
Eu
Gd
Tb
Dy
Ho
Er
Tm
Yb
Rb Ba Th U Nb Ta La Ce Pb Pr Sr P Nd Zr Hf SmEu Ti Gd Tb Dy Y Ho Er TmYb Lu
Lu
1000
1000
b
Upper c ontinental c rust
d
Tholeiitic basalts
Rock Primitive mantle
Rock/Chondrite
OIB
100
E m e is
h a n lo
w - Ti
b a s a lt
s
10
Upper continental crust
100
OIB
10
Emeishan low-Ti b asalts
1
1
La
Ce
Pr
Nd
Sm
Eu
Gd
Tb
Dy
Ho
Er
Tm
Yb
Lu
Rb Ba Th U Nb Ta La Ce Pb Pr Sr P Nd Zr Hf Sm Eu Ti Gd Tb Dy Y Ho Er Tm Yb Lu
Fig. 6. (a, b) Chondrite-normalized REE diagrams. (c, d) Primitive mantle-normalized spidergrams for the Zhulongguan basalts. The normalization values and the ocean-island
basalt (OIB) are from Sun and McDonough (1989). The values of the upper continental crust are from Rudnick and Gao (2003). Data for the Deccan high-Ti picrite is from
Melluso et al. (2006), and the Emeishan low-Ti basalts are from Xu et al. (2001).
significant fractional crystallization or crustal contamination, especially the tholeiitic basalts. The correlations between Ni, V and Cr
(Fig. 8) suggest that the magma primary to the Zhulongguan basalts
might experience varying degree of clinopyroxene- and olivinecontrolled fractionation. For the alkaline basalts, the weak Eu and
Sr anomalies imply minor fractionation crystallization of plagioclase. For tholeiitic basalts, the fractionation of plagioclase explains
the negative Eu (Eu/Eu* = 0.77–0.94) and Sr (Sr/Sr* = 0.29–0.70;
see Niu and O’Hara, 2009) anomalies. Collectively, the Zhulongguan tholeiitic basalts mainly underwent crystal fractionation of
Fig. 7. (A) Sr–Nd isotopic compositions for the Zhulongguan basalts. Plotted for comparison are: the modern depleted upper mantle (N-MORB) (Zimmer et al., 1995), OIB
(White and Duncan, 1996), and EMI and EMII member (Hart, 1988). The CFBs for comparison include low-Ti and high-Ti basalts from Siberian (t = 248 Ma; Sharma et al., 1992),
Emeishan (t = 250 Ma; Xu et al., 2001), and Deccan (t = 66 Ma; Melluso et al., 2006). (B) Nd–Sm isochron diagram for the Zhulongguan basalts. The data for mafic volcanic
rocks on King Island of Tasmania and isochron results are from Meffre et al. (2004); Mt Wright volcanics in western New South Wales from Crawford et al. (1997).
58
X. Xu et al. / Precambrian Research 257 (2015) 47–64
1000
1000
Tholeiitic basalts
ol
Alkaline basalts
ol
cpx
cpx
hb
100
100
V
Ni
10
10
A
1
B
1
1
10
100
1000
Cr
1
10
100
1000
Cr
Fig. 8. (A) Ni and (B) V versus Cr diagrams mainly showing olivine and clinopyroxene fractionation for alkaline and tholeiitic basalts. The vectors are from Li et al. (2010a).
olivine + clinopyroxene + plagioclase, whereas the alkaline basalts
experienced dominantly clinopyroxene and olivine fractionation
to a limited extent.
6.1.2. Crustal contamination
In general, intraplate continental basalts display greater elemental and isotopic diversity than oceanic counterparts, which
have been attributed to varying degrees of interaction between
the continental lithospheric and asthenospheric sources (Wilson,
1989; Arndt and Christensen, 1992; Hawkesworth et al., 1995;
Turner and Hawkesworth, 1995; Xu et al., 2001; Li et al., 2013b;
Wang et al., 2008, 2014).
Nb–Ta and neighboring elements (Th, U and La) are not fractionated from each other during partial melting or fractional
crystallization (Hofmann, 1988), but the enrichment of mantle
source and the crustal contamination can significantly increase LILE
and LREE contents and decrease HFSE/LILE or HFSE/LREE ratios.
For alkaline basalts, the higher Nb/Th (14–25), Nb/U (57–113)
and Nb/La (1.6–2.7) ratios than those of the primitive mantle
(Nb/Th = 8.4; Nb/U = 34; Nb/La = 1.04; Sun and McDonough, 1989)
values reflect the primary signature of the mantle sources without
significant crustal contamination (Fig. 9b and d). The recognition
is also supported by a positive anomaly in Ti (Fig. 6c). Likewise,
their high and positive εNd (T) values (Fig. 7a–c) imply insignificant
crustal contamination for alkaline lavas.
However, some tholeiitic basalts exhibits crust-like characteristics with obvious enrichment in Th, U, LREE and depletion in
Nb, Ta (La/Nb > 1), although a few show no visible HFSE depletion
(La/Nb < 1) ascribing to less/no contamination (Fig. 9b and d). The
contents of Th and U are suggested to be enriched in the upper
continental crust but depleted in the lower continental crust and
lithospheric mantle (Rudnick and Gao, 2003). Therefore the high Th
content (>2.5 ppm) and Th–U positive anomaly indicate contamination with upper crustal materials (Figs. 6d and 9d). Given the
relatively wide range of εNd (T) (−3.1 to −0.2), we consider that the
primary magma must have experienced significant contamination
of upper crustal rocks with low Nd isotopic values although we cannot preclude the assimilation of the metasomatized subcontinental
lithospheric mantle.
Indeed, Fig. 9 shows a general trend toward more crustal contribution from alkaline to tholeiitic basalts. Trace-element ratio–ratio
plots (Fig. 10) for these basalts show good hyperbolic correlations
between Lu/Hf and Hf/Yb, Lu/Hf and Zr/Yb, also indicating crustal
contamination in the form of a binary mixing (Wang et al., 2008).
However these appearances could not been interpreted as the
simple comagmatic evolution with the AFC process on account
of the enrichment of LREE (Fig. 9c). Hence, it seems impossible
that the tholeiitic rocks are derived directly from the alkali lavas.
This conclusion is consistent with the different Nd isotopic values
in the alkaline (+4 to +5) and some of tholeiitic rocks (+2 to +3)
without obvious assimilations. In summary, the magma primitive
to the tholeiite basalts is more depleted in trace elements and
derived from a more depleted source than the magma primitive to
the alkaline basalts.
6.1.3. Tectonic setting
Mantle source compositions and melting conditions determine
the compositions of the basaltic magmas (Cox, 1980; Xu et al.,
2001; Niu and O’Hara, 2008; Li et al., 2013b; and reference herein).
The enrichment in HFSE and LREE of the alkaline basalts may
be directly derived from the asthenospheric mantle such as the
OIB-like source or small degree partial melting of a normal-type
MORB source; but the lower εNd (T) (+4 to +5) values than that of
the contemporaneous depleted upper mantle (εNd (600 Ma) = +8.7)
precludes the latter. Of the alkaline basalts, 11QL-65 is the leastevolved with the highest MgO (11.1 wt.%), Mg-number (68.7) and
compatible element content (Cr = 623 ppm; Ni = 280 ppm). We calculated the major element composition of the primary magma
for this sample according to the procedure of Lee et al. (2009).
Because of the MgO content (>9%), the low pressure fractionation is corrected by incrementally adding olivine. The final primary
magma contains ∼50.9% SiO2 , ∼16.4% MgO, and ∼10.8% FeOt, which
is a picritic composition and corresponds to a melt temperature
of ∼1448 ◦ C (under anhydrous melting condition). The potential temperature (Tp = 1493 ◦ C) of the mantle source is obtained
in terms of the equation of Tp (◦ C) = 1463 + 12.74MgO–2924/MgO
(Herzberg and O’Hara, 2002). The Tp is obviously higher than that
of the modern mid-ocean ridge basalts (1280–1400 ◦ C) and close
to that of the Hawiian picrites (1500–1600 ◦ C) (Putirka, 2005;
Putirka et al., 2007; Herzberg et al., 2007; Herzberg and Asimow,
2008; Lee et al., 2009 and reference herein), indicating an anomalously hot mantle source. Extremely high La/Sm (2.1–2.5) and
Sm/Yb (2.2–3.2) ratios may suggest that they originated from the
garnet-bearing mantle reservoir and experienced the low degree
of partial melting (e.g. Niu et al., 2011). As a result, the primary
magma of the alkaline suite is possibly generated from the partial melting of the asthenospheric mantle caused by a mantle
plume.
For tholeiite basalts, the relatively low degree of fractionation
between HREE and LREE may imply a higher degree of melting
and shallower source than the lavas of the alkaline suite (Niu
et al., 2011). Overall, the ratios of Zr/Y (4–6.2) and Zr/Sm (21–34)
X. Xu et al. / Precambrian Research 257 (2015) 47–64
10
59
7
Tholeiitic b asalts
Alkaline basalts
FC
Primary mantle
a
b
5
5
εNd (T)
εNd (T)
3
FC
0
AFC
1
u
Cr
-1
-5
sta
o
lc
nt
am
a
in
tio
n
-3
-10
-5
0
6
4
2
8
10
0.5
0
12
1.0
MgO
1.5
2.0
2.5
7
30
c
d
25
3
20
Nb/Th
5
Cr
1
us
ta
lc
εNd (T)
3.0
Nb/La
Cr
10
on
-1
Primary m antle
15
ta
m
in
on
ta
ati
on
Upper c rust
5
at
-3
us
c
tal
n
mi
io
n
0
-5
0
1
3
2
5
4
0.5
0
1.0
1.5
2.0
2.5
3.0
Nb/La
La/Sm
Fig. 9. Plots of (a) MgO versus ␧Nd (T); (b) Nb/La versus ␧Nd (T); (c) La/Sm versus ␧Nd (T); (d) Nb/La versus Nb/Th. The ratios of the primary mantle (Sun and McDonough,
1989) and upper continental crust (Rudnick and Gao, 2003) are also plotted for comparison.
are similar to many intra-plate basalts (Zr/Y > 3.5, Zr/Sm ≈ 30), but
are distinct from those of island arc rocks (Zr/Y < 3.5, Zr/Sm < 20)
(Fig. 11a; Wilson, 1989). All tholeiitic and alkaline basalts are
dropped into continental flood basalts or ocean-island and alkaline basalts field in the Ti–V diagram (Fig. 11b). In addition, the
clinopyroxene phenocrysts of the Zhulongguan basalts show a clear
rifted-related trend (Fig. 11c).
The coexistence of high-Ti and low-Ti groups is recognized
widely in continental flood basalt provinces, such as the Parana and
the Karoo (Gibson et al., 1995), Deccan Traps (Melluso et al., 2006),
7.0
2.0
Tholeiitic basalts
A
B
Alkaline b asalts
1.8
6.5
6.0
R² = 0 .9878
1.4
Zr/Y
Hf/Yb
1.6
1.2
5.0
1.0
4.5
0.8
4.0
0.6
0.05
0.10
0.15
Lu/Hf
0.20
R² = 0 .9397
5.5
3.5
0.05
0.10
0.15
0.20
Lu/Hf
Fig. 10. Trace-element ratio-ratios plots of Hf/Yb and Zr/Y against Lu/Hf for the Zhulongguan basalts. The linear or hyperbolic curves are reflective of binary mixing or
contamination with crust materials (Niu and Batiza, 1997).
60
X. Xu et al. / Precambrian Research 257 (2015) 47–64
Fig. 11. Discrimination plots for the Zhulongguan basalts. (a) Zr versus Zr/Y diagram after Pearce and Norry (1979); (b) Ti versus V plot after Shervais (1982). The fields of arc
tholeiitic, MORB, continental flood basalts, ocean-island and alkali basalts were drawn by Rollinson (1993) according to Shervais (1982); (c) Alz (percentage of tetrahedral
sites occupied by Al) versus TiO2 in clinopyroxenes (Loucks, 1990) for the Zhulongguan basalts.
Siberian Traps (Hawkesworth et al., 1995; Sharma et al., 1992),
Emeishan (Xu et al., 2001), South China (Wang et al., 2008), and
North Qaidam (Song et al., 2010). In conclusion, the Zhulongguan
tholeiitic (low-Ti) and alkaline (high-Ti) basalts are likely formed in
an extension-related within-plate environment probably induced
by a mantle plume, rather than the supra-subduction zone.
6.2. Continental rifting and Qilian-Ocean formation
The formation history of the Qilian Ocean has not been well constrained so far, due to the lack of the bimodal volcanics or intraplate
magmatism that could indicate the break-up phase. Based on the
anorogenic granitic intrusions with ages of 750–800 Ma (Tseng
et al., 2006; Tung et al., 2013) and relic cores in zircons from some
eclogites (Zhang et al., 2007), Song et al. (2013) inferred that the
seafloor spreading of the Paleo-Qilian Ocean, as separating South
China, Qilian-Qaidam and Tarim blocks from Rodinia, might started
at least from ∼710 Ma and closed at ∼445 Ma.
This study is the first report on the presence of ∼600–580 Ma
rift volcanism in the north margin of the Qilian-Qaidam block, NW
China. The association of alkaline-tholeiitic lavas with shallowmarine facies sedimentary layers in the Zhulongguan Group is
strongly akin to an extensive mafic volcanic passive margin prior
to the continental break-up, e.g. the seaward-dipping reflector
sequences (SDRS) in southeastern Australia (Direen and Crawford,
2003; Meffre et al., 2004). The similar isochron results in Fig. 7B
demonstrate that the Zhulongguan basalts resemble the volcanic
rocks in Australia at the end of the Precambrian. The oldest
550 Ma ophiolite in the North Qilian Orogen (Shi et al., 2004;
Song et al., 2013) and the Marlborough terrane of the northern
New England Fold Belt (Bruce et al., 2000) may also support that
the opening of the Paleo-Qilian Ocean (a branch of Proto-Pacific
Ocean) may occurred immediately after the ca. 600–580 Ma rifting
event.
6.3. Implications for the palaeogeographic position of the
Qilian-Qaidam block in Rodinia
In the last decades, much attention had been paid to the geological evolution of the Rodinia supercontinent. However, its breakup
history remains controversial because of the continuity and complexity of widespread rifting associated with episodic plume pulses
(Hoffman, 1991; Powell et al., 1993, 1994; Veevers et al., 1997;
Preiss, 2000; Wang and Li, 2003; Li et al., 1999, 2003, 2008b;
Cawood, 2005; Ernst et al., 2008). It had been proposed that the
early rift events (820–750 Ma) that widely occurred in the rim of
Proto-Pacific Ocean were followed by the initial breakup of Rodinia
(Powell et al., 1994; Park et al., 1995; Wingate et al., 1998; Li et al.,
1995, 1999, 2003, 2008b; Meffre et al., 2004; Cawood, 2005). Similarly, anorogenic magmatism during this time interval has been
also reported in the Qilian and Qaidam region (Li et al., 2005; Tseng
et al., 2006; Lu et al., 2008; Chen et al., 2009b; Song et al., 2010;
Tung et al., 2013). Combined with the comparable Grenville-age
orogeny, the recognition demonstrates the strong affinity between
the Qilian-Qaidam (including Quanji Massif) and South China, even
Tarim blocks (Wan et al., 2001, 2006; Lu et al., 2008; Chen et al.,
2009b; Tung et al., 2007, 2013; Song et al., 2010, 2012, 2014).
The relatively late stage of intraplate magmatic events
(∼600 Ma) have been documented in Australia, Tarim, and South
China (Table 4). This may represent a long duration fragmentation of Rodinia and the waning stage of plume volcanism (Xu et al.,
2013). In north margin of Tarim (present orientation), ca. 615 Ma
layered basalts are thought to have been generated in an intracontinental rift environment (Xu et al., 2009, 2013). During late
Neoproterozoic to Early Paleozoic, South China received deposition
of thick, platform-type carbonates, phosphorite and black shales by
rapid subsidence (Wang and Li, 2003), but without lava flows (Shu
et al., 2011). Only two volcanic ash beds with ages of 621 ± 7 Ma and
555 ± 6 Ma were recognized in the terminal Proterozoic Doushantuo Formation (Zhang et al., 2005), which means that the South
China was far away from the volcanic eruption centers.
In western New South Wales of Australia, the transitional alkaline basalt-rhyolite suite with zircon SHRIMP age of 586 ± 7 Ma
occurs together with marine sedimentary rocks, which has been
demonstrated to represent a continental rift setting (Crawford et al.,
1997). On King Island, a thick (>900 m) sequence of late Neoproterozoic volcanic and intrusive rocks plus shallow marine carbonates
and siltstones was formed; the upper tholeiitic basalts and picrite
gave a Sm–Nd isochron age of 579 ± 16 Ma (Meffre et al., 2004)
(Fig. 7b). In western Tasmania, a typical rift succession of tholeiitic basalts plus shallow water carbonates, including the Rocky
Cape dyke swarm (590 ± 8 Ma), suggests a latest Precambrian passive continental margin (Crawford, 1992). Direen and Crawford
(2003) concluded that the late Neoproterozoic (600–580 Ma) passive margin volcanic rocks occurred along three elongate belts in
the Delamerian Orogen of southeastern Australia, the Wonominta
Block of western New South Wales, the Gleneleg Zone of western Victoria and the King Island-western Tasmania, respectively. In
addition, the phase of intra-basin fluid flow with a high 87 Sr/86 Sr
value of 0.7180 in the Adelaide Geosyncline was terminated at
∼586 Ma, which may indicate the onset of a new phase of extension in Australian eastern margin (Foden et al., 2001). Furthermore,
mafic schists associated with psammitic rocks at ca. 600 Ma from
the Anakie Inlier in northeastern Australia were also suggested to
X. Xu et al. / Precambrian Research 257 (2015) 47–64
61
Table 4
Compilation of intraplate magmatic records during the late Neoproterozoic in east margin of proto-Gondwana.
Location
Southeast Australia
Mt Wright and Arrowsmith
Western Tasmania
Tasmania (King Island)
Western Victoria and South Australia
Qilian-Qaidam
The western segment of North Qilian
South China
Yangtze Gorge (Doushantuo Formation)
Tarim
WE Tarim (Aksu area)
NE Tarim (Quruqtagh area)
Rock association
Age and method
Tectonic setting
Reference
Alkaline basalt-rhyolite
and interbeded marine
sedimentary
Picrite, alkaline to tholeiitic
basalts, dolerite dykes,
dolomitic limestone
Picrite lavas, tholeiite,
siliciclastics and
carbonates
Tholeiitic to transitional
alkaline basalts,
limestones, volcaniclastic
and picritic lavas
586 ± 7 Ma (SHRIMP)
Continental rift
Crawford et al. (1997)
588 ± 8 Ma and 600 ± 8 Ma
(K-Ar dates)
Crawford and Berry (1992)
579 ± 16 Ma (Sm–Nd isochron
age)
An attenuated, rifted
passive continental
margin
Passive continental
margin
Ca. 589–501 Ma
Rift
Direen and Crawford (2003)
Alkaline and tholeiitic
basalts, volcaniclastic,
dolomitic limestone,
sandstone, siltstone,
iron-ore layer
600–583 Ma (Zircon SIMS)
Continental rift
This study
Volcanic ash, shale,
limestone
555.2 ± 6.1 Ma and 621 ± 7 Ma
(SIMS)
?
Zhang et al. (2005)
Transitional basalts,
sandstone, siltstone,
dolostone
Basaltic and andesitic
lavas, pyroclastic, siltstone
and sandstone
615.2 ± 4.8 Ma, 614.4 ± 9.1 Ma
(SHRIMP)
The waning stage of
plume
Xu et al. (2013)
615 ± 6 Ma (SHRIMP)
Related to the Rodinian
breakup
Xu et al. (2009)
be developed in a passive continental margin setting (Fergusson
et al., 2009).
These magmatism during the latest Neoproterozoic described
above apparently record the simultaneous rifting on the eastern
margin of Australia–Antarctica and imply the uniform affinities
among these blocks before continental drifting. It is worth noting
that the rifting-related records slightly predates the second continental breakup around 550 Ma in the periphery of the Proto-Pacific
Ocean (Bond et al., 1984, 1985; Meert et al., 1994; Powell et al.,
1994; Veevers et al., 1997). The Yushigou ophiolite (550 Ma) of
the North Qilian and the Marlborough terrane (562 Ma) of eastern Australia may represent the initial opening of oceanic basin
(Bruce et al., 2000; Shi et al., 2004; Song et al., 2013). Overall,
renewed rifting related to the opening of the marginal sea could
be indicated by the 615–580 Ma magmatism occurred in east margin of proto-Gondwana (Table 4), which may record the separation
of the Chinese blocks from there. The protracted history of rifting can be interpreted by the process of episodically separating
of microcontinents such as the Qilian-Qaidam, which faced the
wide proto-Pacific Ocean (Powell et al., 1994; Direen and Crawford,
2003; Fergusson et al., 2009).
Fig. 12 illustrates the configuration of the eastern protoGondwana at ca. 600–580 Ma. The palaeogeographic positions of
South China, East Antarctica and Laurentia are in general consistent with the proposal by Li et al. (1999, 2003, 2008b), although
other configuration models about Tarim and South China have been
also suggested (Lu et al., 2008; Li et al., 2013a,b,c; Zhang et al.,
2013a,b; Yao et al., 2014). The original unity between Australia,
South China, Qilian-Qaidam, and Tarim blocks was suggested by
the plume-related radial dyke swarms (∼825 Ma) (Lu et al., 2008).
Until 600 Ma South China was far from both Laurentia and Australia
given by the lack of rift magmatism. At that time, the Qilian-Qaidam
block, as a connection between Australian and South China, began
rifting and dispersal into the Proto-Pacific Ocean. Considering the
Meffre et al. (2004)
compression of the Pan-African orogeny (500–700 Ma) in the interior of Gondwanaland (Hoffman, 1991; Veevers, 2003; Li et al.,
2008b), it seems appropriate to put these blocks at the external
margin of continent in the extensional regime. The simultaneous
assembly and breakup of different parts of a supercontinent are not
only recognized in the cycle of the Gondwana, but also exist in the
evolution of the Rodinia supercontinent (Ernst et al., 2008).
600-550 Ma
QR
Australia
AI
?
Tarim
AK
WA
Proto-Pacific
Ocean
GZ
East
Antarctica
KT
NQ
Qilian-Qaidam
South China
Passive continental margin
Middle oceanic ridge
Rifting-related volcanic-sedimentary
sequence (ca. 600-580 Ma)
Fig. 12. Reconstruction of the East Gondwana at ca. 600–580 Ma (modified after
Fergusson et al., 2009). Relative positions of Australia, East Antarctica, and South
China are based on the proposal by Li et al. (1999, 2003, 2008b). The Tarim had been
tentatively placed in close to the eastern Australia (Lu et al., 2008). The 600–580 Ma
passive margin volcanics include: WA = Mt Wright and Arrowsmith in western New
South Wales; GZ = Glenelg Zone in western Victoria; KT = King Island-Tasmania;
NQ = North Qilian; AI = Anakie Inlier in northeastern Australia; AK = Aksu area in WE
Tarim; QR = Quruqtagh area in NE Tarim (also see Table 4).
62
X. Xu et al. / Precambrian Research 257 (2015) 47–64
Acknowledgements
We thank Xianhua Li and his laboratory group for helping
with SIMS dating, G.Z. Li and W.P. Zhu for Sr–Nd isotopic analyses. We also thank Editors and Reviewers for their constructive
comments. This study was supported by the Major State Basic
Research Development Program (2015CB856105), Basic geological survey program of China Geological Survey (1212011121258)
and National Natural Science Foundation of China (Grant Nos.
41372060, 40825007, 41121062, 41130314).
Appendix A. Supplementary data
Supplementary data associated with this article can be found,
in the online version, at http://dx.doi.org/10.1016/j.precamres.
2014.11.017.
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