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Geochimica et Cosmochimica Acta 143 (2014) 189–206
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Origin of carbonatites in the South Qinling orogen:
Implications for crustal recycling and timing of collision
between the South and North China Blocks
Cheng Xu a,⇑, Anton R. Chakhmouradian b, Rex N. Taylor c, Jindrich Kynicky d,
Wenbo Li a, Wenlei Song a, Ian R. Fletcher e
a
Laboratory of Orogenic Belts and Crustal Evolution, School of Earth and Space Sciences, Peking University, China
b
Department of Geological Sciences, University of Manitoba, Canada
c
School of Ocean and Earth Science, University of Southampton, UK
d
Department of Geology and Pedology, Mendel University, Brno, Czech Republic
e
Department of Applied Geology, Curtin University, Australia
Available online 13 April 2014
Abstract
Most studies of compositional heterogeneities in the mantle, related to recycling of crustal sediments or delaminated subcontinental lithosphere, come from oceanic setting basalts. In this work, we present direct geochronological and geochemical
evidence for the incorporation of recycled crustal materials in collision-related carbonatites of the South Qinling orogenic belt
(SQ), which merges with the Lesser Qinling orogen (LQ) to separate the South and North China Blocks. The SQ carbonatites
occur mainly as stock associated with syenites. The data presented here show that zircon from the syenites yields an age of
766 ± 25 Ma, which differs significantly from the age of primary monazite from the carbonatites (233.6 ± 1.7 Ma). The syenites contain lower initial 87Sr/86Sr and higher eNd values. This indicates that the carbonatites do not have genetically related
with the silicate rocks, and were directly derived from a primary carbonate magma generated in the mantle. The carbonatites
show a Sr–Nd isotopic signature similar to that of the chondritic uniform reservoir (CHUR), and parallel Sm–Nd model ages
(TCHUR) of 190–300 Ma. However, the rocks have extremely variable Pb isotopic values straddling between the HIMU and
EM1 mantle end-members. Most carbon and oxygen isotopic compositions of the SQ carbonatites plot outside the field for
primary igneous carbonates. Their d13C shows higher value than a ‘normal’ mantle, which implies an incorporation of recycled inorganic carbon. The carbonatites were emplaced close to the Mianlue suture, and followed the closure of the Mianlue
ocean and Triassic collision of the South and North China Blocks. However, direct melting of the subducted Mianlue oceanic
crust characterized by high eNd and low (EM1-like) 206Pb/204Pb values cannot explain the CHUR-like Nd signature and the
Pb isotopic trend toward HIMU in the SQ carbonatites. We conclude that their parental magma was derived from a source
incorporating the Mianlue oceanic crust mixed with an asthenospheric (or deeper) material characterized by high Pb and low
Nd isotopic values. This material represents a deep-seated Proterozoic carbonate component recycled via mantle convection
or localized upwelling. Notably, this model cannot explain the isotopic compositions of the Late Triassic (209–221 Ma) carbonatites in the LQ, characterized by a mantle-derived d13C, but EM1-like Sr–Nd–Pb isotopic compositions. This signature is
best explained in terms of delamination of the lower continental crust thickened during the collision of the South and North
China Blocks, and partial incorporation of the delaminated material into the LQ mantle source. Modeling of the measured
Sr–Nd–Pb isotopic variations suggests that the source of the LQ carbonatites could be produced by mixing of 80–85% of
⇑ Corresponding author. Tel.: +86 1062753894.
E-mail address: [email protected] (C. Xu).
http://dx.doi.org/10.1016/j.gca.2014.03.041
0016-7037/Ó 2014 Elsevier Ltd. All rights reserved.
190
C. Xu et al. / Geochimica et Cosmochimica Acta 143 (2014) 189–206
mantle material and 15–20% of delaminated lower continental crust. The emplacement of the SQ and LQ carbonatites marked
a gradual transition from a compressional tectonic regime, brought about by the collision of the South and North China
Blocks to intra-orogenic extension in the waning stages of the Triassic orogeny.
Ó 2014 Elsevier Ltd. All rights reserved.
1. INTRODUCTION
2. GEOLOGICAL SETTING AND PETROGRAPHY
Carbonatites are mantle-derived carbonate-rich igneous
rocks found in extensional settings. Most of the presently
known carbonatite occurrences are situated in anorogenic
settings related to intracontinental rifting. Collision-related
carbonatites are much rarer (Tilton et al., 1998; Hou et al.,
2006; Chakhmouradian et al., 2008) and remain poorly
understood. These rocks provide valuable information on
the composition of the mantle because (1) their isotopic
characteristics are generally inherited from the mantle
source owing to the initially high concentrations of Sr
and rare-earth elements (REE) in carbonatitic magmas
(e.g., Nelson et al., 1988), and (2) carbonate melts have
extremely low viscosities (Treiman, 1989) facilitating their
rapid ascent to the surface. The importance of recycled sediments and other crustal materials as a source for carbonatitic magmatism is disputed (Barker, 1996; Hoernle et al.,
2002; Halama et al., 2008; Bell and Simonetti, 2010). Previously published isotopic studies of young carbonatites
(<200 Ma) identified a high-l material (HIMU), Enriched
Mantle 1 (EM1) and Focal Zone (FOZO) as the major
mantle components involved in the generation of carbonatitic magmas (Tilton and Bell, 1994; Simonetti et al., 1998;
Bell and Tilton, 2001, 2002). At present, the prevailing view
is that these magmas are derived from a lithospheric source
affected by asthenospheric upwelling or plumes involving
deeper parts of the mantle (e.g., Bell and Tilton, 2001;
Tolstikhin et al., 2002). This conventional isotopic framework is not applicable to recently discovered young carbonatites in the Qinling orogenic belt of central China. The
Qinling belt is the product of ocean closure and continental
collision in the Triassic, and hosts several carbonatite intrusions of Triassic age. These rocks provide critical information about deep mantle sources in the collisional
environment and the role of crust recycling in magma generation. Previous studies show that the carbonatites not
associated with alkalic silicate rocks from the northernmost
Qinling have less radiogenic Nd values than EM1 (Huang
et al., 2009; Xu et al., 2011), and do not record simple mixing between the HIMU and Enriched Mantle components.
However, the southernmost Qinling carbonatites examined
in this study are associated with syenites and characterized
by slightly depleted isotopic compositions. Their origin and
genetic relationship with the syenites are unknown. Thus,
our primary objective is to develop alternative models of
carbonatite petrogenesis in collisional settings that will
explain both the structural relations and unusual isotopic
signatures of the Qinling rocks. In a broader context, correct interpretation of the geodynamic evolution of the Qinling belt gleaned from isotopic data is important in
unraveling the tectonic history of East Asia.
The Qinling zone is linked to the Dabie-Sulu ultrahigh-pressure metamorphic belt to the east (Fig. 1) and separates
the North China Block from the South China Block. Closure
of the Mianlue Ocean in the Triassic resulted in collision of
the South China Block with the South Qinling belt (SQ)
along the Mianlue suture, followed by amalgamation of
the South and North China Blocks (Meng and Zhang,
1999; Ratschbacher et al., 2003). The southern margin of
the North China Block was deformed in the process to produce the Lesser Qinling belt (LQ), which, together with the
North Qinling belt (NQ), forms a passive continental margin
(Xue et al., 1996). The Qinling orogen is bounded by the
Lingbao–Lushan–Wuyang fault in the north and the Mianlue–Bashan–Xiangguang fault in the south. Two suture
zones incorporating ophiolitic mélange, the Shangdan suture
in the north and the Mianlue suture in the south, are well documented (Meng and Zhang, 1999; Ratschbacher et al., 2003).
The orogen is divided into the aforementioned SQ, NQ and
LQ terranes, which are separated by the two sutures and several major faults (Fig. 1). The LQ is characterized by highly
deformed, high-grade metamorphic basement rocks of the
Mesoarchean Taihua group and granite-greenstone assemblages of the Neoarchean Dengfeng group (Kröner et al.,
1988). The carbonatites intruded the Taihua group, which
comprises amphibolite- to granulite-facies metamorphic
rocks discordantly overlain by Mesoproterozoic volcanics.
In the SQ, the Proterozoic basement consists of the Paleoproterozoic Douling group comprising paragneiss and granitic
gneiss, Paleoproterozoic- to Neoproterozoic Wudang and
Yunxi groups composed of alkali basalt, alkali trachyte,
dacite, rhyolite and pyroclastic rocks, and the Neoproterozoic Yaolinghe group made up of tholeiitic basalt, spilitic diabase, alkali trachyte and spilite (Huang, 1993).
Many occurrences of carbonatites have now been identified in this part of China; these rocks were emplaced into
the deformed southern margin of the North China Block
in the LQ and into the northern margin of the South China
Block in the SQ. The SQ carbonatites are tectonically confined to the southern margin of the Mianlue suture. They
were emplaced into syenitic rocks as stock and minor dikes,
the largest of which covers 0.3 km2 at the current erosion
level. The syenites were altered and impregnated by numerous calcite-bearing veinlets. The rock is composed predominantly of alkali feldspar and plagioclase (Fig. 2a); minor
and accessory phases include biotite, amphibole, clinopyroxene, quartz, sulfides, calcite, fluorapatite, barite, REE
minerals, rutile and zircon. The SQ carbonatites are
composed predominantly of medium- to fine-grained
calcite. Minor and accessory phases include biotite,
feldspar, fluorapatite, ilmenite, pyrite, barite, fluorite,
C. Xu et al. / Geochimica et Cosmochimica Acta 143 (2014) 189–206
191
Fig. 1. The location and geological setting of the carbonatites in the South Qinling (SQ) and Lesser Qinling orogenic belts (LQ-S in Shanxi,
LQ-H in Henan; modified after Xu et al., 2010b, 2011).
quartz, monazite, bastnäsite, synchysite, ferrocolumbite
and rutile. Monazite is an early-crystallising mineral in
the carbonatites and occurs as ovoid grains up to 200 lm
across. The crystal does not exhibit patchy zoning
(Fig. 2c), and contains little compositional variation and
constant (La/Nd)CN ratio of 3 (Xu et al., 2010a). The ages
of the SQ carbonatites and syenites have not been well constrained prior to the present work. In the LQ, the carbonatites are situated in the central part of the Shanxi Province
(LQ-S) and in the western part of the Henan Province (LQH). They occur as dikes containing abundant sulfide mineralization including molybdenite, and are not associated
with alkalic silicate rocks. Rhenium-Os measurements on
molybdenite from LQ-S and LQ-H samples gave ages of
221.5 Ma (Du et al., 2004) and 209.5 Ma (Huang et al.,
2009), respectively. The detailed geological and petrological
information on LQ carbonatites was reported by Xu et al.
(2007, 2011) and Huang et al. (2009).
3. ANALYTICAL METHODS
Because the SQ carbonatite lacks zircon or baddeleyite,
monazite was chosen for a radiochronological study of the
rock. Its age was determined by U–Th–Pb methods using a
sensitive high-resolution ion microprobe (SHRIMP) at the
Curtin University, Australia. The mineral was separated
from fresh samples using standard techniques, cast into
25 mm resin mounts together with grains of 425-Ma reference material 44069 (Aleinikoff et al., 2006) and polished
to expose flat sections. The primary beam diameter was
25 lm, and primary O2 beam current was 1.5 nA. A mass
resolution (M/DM at 1% of peak height) was 5000, and
each analysis consisted of six scans of the mass spectrum.
Data were acquired in a single analytical session, following
the procedures described in Foster et al. (2000) and Fletcher
et al. (2010). The primary data ratios were 206Pb+/270[UO2]+
and 208Pb+/264[ThO2]+, both of which were subjected to
1-dimensional calibrations (Fletcher et al., 2010) with
standard 44069 as the primary age reference. Additional
standard materials French, z2234 and 2908 provided reference U, Th and Y concentrations. Corrections for matrix
effects in Pb/U and Pb/Th were made subsequently, in
spreadsheet templates, using matrix correction factors from
Fletcher et al. (2010). Data from French, z2234 and z2908
provided first-order confirmation of the correction factors
for this session.
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C. Xu et al. / Geochimica et Cosmochimica Acta 143 (2014) 189–206
a
b
c
Fig. 2. Photomicrographs of the South Qinling syenites (a) and carbonatites (b, c). Kfs, K-feldspar; Ab, albite; Aug, augite; Bt, biotite; Ap,
fluorapatite; Cal, calcite; Mn, monazite.
Zircon crystals were extracted from representative syenite samples using conventional magnetic and heavy-liquid
separation techniques, and then mounted in epoxy and polished. The crystals were examined using cathodoluminescence (CL) imaging to reveal their internal structure. Age
determination was done using laser-ablation inductivelycoupled-plasma mass-spectrometry (LA-ICPMS). For
LA-ICPMS measurements, we used an Agilent 7700
mass-spectrometer coupled to a 193-nm ArF excimer laser
at the Institute of Geochemistry, Chinese Academy of Sciences (CAS). The laser beam was focused on the sample
with a fluency of 10 J/cm2 and a spot of 32 lm in diameter
at a repetition rate of 5 Hz for 40 s. Helium was used as a
carrier gas to transport the ablated aerosol to the massspectrometer. Zircon Nancy 91500 was used as an external
calibration standard to correct for instrumental mass bias
and elemental fractionation. Zircon standards GJ-1 and
Plešovice were employed for quality control. The Pb content of zircon was externally calibrated against NIST 610
with Si as an internal standard, whereas other trace elements were measured with Zr as an internal standard.
Raw-data reduction was performed off-line using the
ICPMSDataCal software (Liu et al., 2010).
The trace-element composition of syenite samples was
determined by solution ICPMS (ELAN DRC-e) at the
Institute of Geochemistry, CAS. The analytical protocol
for trace elements was described in detail in Liang et al.
(2000). Replicate analyses (including well-characterised
standards) indicated that the accuracy of trace element
measurements was better than 10%. The Rb, Sr, Sm, Nd,
U, Th and Pb concentrations of carbonatites and their constituent minerals were analyzed by solution ICPMS
(Thermo Fisher Scientific X-Series II) in the School of
Ocean and Earth Science at the University of Southampton
(UK). The instrument was calibrated using five international rock standards, BIR-1, JGb-1, JB-3, BHVO-2 and
JB-1a, which were run before, during, and after each group
of analyses. The data processing procedure included linear
drift correction, internal (matrix) correction, REE and Ba
interference corrections, blank subtraction, calibration with
international standards, and a dilution correction. In-run
precision was routinely in the 1–5% range, and accuracy relative to the reference materials 5%. In situ LA-ICPMS
determinations of Rb, Sr, U, Th and Pb concentrations of
calcite were performed in polished sections at the Australian National University (ANU) using an analytical methodology similar to that described below for the Pb
isotopic measurements. The calculation of element concentrations was like that reported by Eggins et al. (1997), and
detection limits were determined using the approach of
Longerich et al. (1996). Analytical precision was <5% at
the ppm level.
Carbon and oxygen isotopic compositions of calcite
from the SQ carbonatites and associated veinlets in the syenites were measured at the Institute of Geochemistry, CAS,
using a continuous-flow isotope ratio mass spectrometer
(IsoPrime). The results were expressed in per mil (&) using
the conventional d-notation, i.e., d = (R1/R2 1) 1000,
where R1 is the 13C/12C or 18O/16O ratio in the sample
and R2 is the corresponding isotope ratios in the standards
C. Xu et al. / Geochimica et Cosmochimica Acta 143 (2014) 189–206
(V-PDB for C and V-SMOW for O). Reproducibilities of
d13C and d18O values, measured as one standard deviation
for the carefully selected reference materials (GBW04405,
GBW04406, GBW04417), were 0.15& and 0.2&, respectively. The Sr and Nd isotopic compositions of the selected
samples from the SQ carbonatites were analyzed by a
seven-collector VG Sector 54 mass-spectrometer with a separable-filament source in the School of Ocean and Earth
Science, University of Southampton. Strontium and Nd
isotopic ratios were determined as the average of >100
ratios by measuring ion intensities in multidynamic collection mode. Mass fractionation corrections for Sr and Nd
isotopic ratios were normalized to 86Sr/88Sr = 0.1194 and
146
Nd/144Nd = 0.7219, respectively. Total blanks were
<1.1 ng for Sr and <<0.2 ng for Nd, which represent
<0.05% of the mass of both elements in the measured fraction. Repeated measurements for the Nd standard JMC321
and Sr standard NBS987 yielded 143Nd/144Nd =
0.511125 ± 11 (2r, n = 45) and 87Sr/86Sr = 0.710252 ± 11
(2r, n = 169), respectively. These values are similar to the
commonly accepted reference values (e.g., Jochum et al.,
2005).
Strontium isotopic composition of calcites from the SQ
carbonatites and veinlets in syenites was analyzed by LAMC-ICPMS using a Thermo Neptune instrument at the
ANU. Two calcite crystals from the same hand specimen
were embedded in a polished epoxy mount with their cores
exposed. The spot diameter used for the measurement was
178 lm for the calcite and 233 lm for the Tridacna standard. The average 87Sr/86Sr ratio obtained for the standard
was 0.70913, i.e., close to the accepted in-house value of
0.70917.
Lead fractions were separated and purified using
HBr and anion exchange resin and then analyzed by
MC-ICPMS (Thermo Neptune) in the School of Ocean and
Earth Science at the University of Southampton. Doublespiking (Southampton–Brest–Lead 207–204 spike, SBL-74)
was used to correct for instrumental mass bias on the
measurements. The Pb isotopic composition of the samples
was obtained from the natural and mixture runs by iterative
calculation adopting an exponential mass-bias correction.
Repeated measurements for Pb standard NBS981 gave
206
207
Pb/204Pb = 16.9404 ± 25,
Pb/204Pb = 15.4973 ± 20
and 208Pb/204Pb = 36.7173 ± 65 (2r, n = 32). Total procedural blanks were Pb < 50 pg, which is considered negligible
relative to the sample loads and analyzed concentrations
(>300 ng Pb). The Pb isotopic composition of calcites from
the SQ carbonatites was also measured in polished thin sections using a HP7500 Agilent LA-ICPMS at the ANU. The
diameter of the ablation spot was 54 lm. The NIST 610
glass was used as a calibration standard for all samples,
with 43Ca as an internal standard for quantitative analysis.
In-run signal intensities for 206Pb, 207Pb and 208Pb were
monitored during the measurement to make sure that the
laser beam stayed within the grain chosen for analysis
and did not penetrate inclusions. The 204Pb content could
not be quantified because of systematic Hg contamination.
The 207Pb/206Pb and 208Pb/206Pb ratios in the calcites were
calculated so that Pb concentration ratios estimated from
206
Pb, 207Pb and 208Pb measurements were multiples of
193
0.9098 (207Pb/206Pb value in the NIST 610 calibration standard) and 2.169 (208Pb/206Pb value in NIST 610: Jochum
et al., 2005), respectively.
4. RESULTS
The SHRIMP and LA-ICPMS U–Th–Pb data for monazite and zircon grains from the SQ carbonatites and syenites, respectively, are given in Tables 1 and 2 and shown
in Fig. 3. Trace elements and C–O–Sr–Nd–Pb isotopic values for the examined syenites and carbonatites are listed in
Tables 3–6 and summarized in Figs. 4–6.
4.1. U–Th–Pb geochronology
Monazite is widely used as a U–Th–Pb geochronometer
to determine the magmatism timing (e.g., Harrison et al.,
1995). Experiment shows that a 10 lm-sized monazite grain
would have a Pb closure temperature in excess of 900 °C
(Cherniak et al., 2004). Carbonatite magmas contain abundant volatiles, which result in low crystallization temperature (Jago and Gittins, 1991). The carbonatite lava flows
in the crater of the volcano Oldoinya Lengai (Tanzania)
show extremely low eruption temperature of 544–590 °C
(Jago and Gittins, 1991). Therefore, the U–Th–Pb system
in monazite is a valuable chronometer for dating carbonatite magmatism. The common Pb contents of the analyzed
monazite samples from the SQ carbonatites are quite variable; all analyses with >2% common Pb (either 208Pb or
206
Pb) were excluded from further consideration, even
though these discarded data are generally consistent with
the retained analyses, implying that the chosen 2% cut-off
limit is sufficiently conservative. Only three 208Pb/232Th
measurements gave “high” common Pb (2.2–3.3%) and
were rejected. Using the Isoplot software (Ludwig, 2003),
the 16 analyzed grains gave a weighted mean age of
235.5 ± 2.1 Ma with MSWD = 0.91. A similar assessment
of the 206Pb/238U data gave a weighted mean age of
231.5 ± 2.6 Ma with MSWD = 1.2. Since the 208Pb/232Th
and 206Pb/238U data were calibrated independently, they
can be combined, giving an age of 233.6 ± 1.7 Ma with
MSWD = 1.3 for the main phase of monazite crystallization (Fig. 3a). The remaining two monazite grains appear
to have been partially recrystallized between 220 and
190 Ma, but the outliers show too much scatter to indicate
a distinct datable post-emplacement recrystallization
(growth?) episode other than the main 235 Ma result.
The younger monazite contains slightly higher U content
and has a lower average Th/U ratio relative to the primary
generation, possibly indicating a crustal input. According
to the Isoplot results for several monomineralic fractions
(biotite, ferrocolumbite, pyrite, fluorapatite and monazite)
from the SQ carbonatites and whole-rock measurements,
the 206Pb/204Pb vs. 207Pb/204Pb (“Pb–Pb”) isochron gave a
combined age of 262 ± 25 Ma with MSWD = 357
(Fig. 3b). This value is (within error) in general agreement
with the SHRIMP monazite age.
Zircons from the syenites are euhedral oscillatory-zoned
crystals >70 lm in length (Fig. 3c). Twenty-four LAICPMS analyses of the zircon crystals show variable Th
194
C. Xu et al. / Geochimica et Cosmochimica Acta 143 (2014) 189–206
Table 1
SHRIMP U–Pb and Th–Pb data from monazites in South Qinling carbonatites, central China.
Analysis
U
(ppm)
Th
(ppm)
Self-consistent Th–Pb group, in
MY-01.4
173
4236
MY-01.12
145
2175
MY-01.26
76
2789
MY-01.19
99
1776
MY-01.15
87
2895
MY-01.21
249
1524
MY-01.20
43
1082
MY-01.23
111
2795
MY-01.9
101
2049
MY-01.24
135
2010
MY-01.25
64
2839
MY-01.7
94
2064
MY-01.11
136
1007
MY-01.2
114
2396
MY-01.6
39
1392
MY-01.10
49
3636
Th/U
4f206
(%)
206
Pb/
U
±s
238
Pb/232Th sequence
25
1.86
0.0355
15
0.24
0.0367
37
1.63
0.0359
18
2.60
0.0362
33
3.11
0.0349
6.1
0.07
0.0373
25
1.36
0.0370
25
6.97
0.0368
20
1.51
0.0371
15
1.55
0.0373
45
6.13
0.0378
22
2.59
0.0364
7.4
0.33
0.0363
21
1.87
0.0367
36
1.00
0.0365
75
2.84
0.0379
t[206Pb/
238
U]
(Ma)
±s
4g208
(%)
208
Pb/
Th
±s
t[208Pb/
232
Th]
(Ma)
±s
232
208
0.0006
0.0006
0.0007
0.0007
0.0007
0.0006
0.0009
0.0008
0.0007
0.0007
0.0009
0.0007
0.0006
0.0007
0.0009
0.0010
224.6
232.3
227.1
228.9
221.0
236.4
234.1
232.8
234.6
236.3
239.3
230.4
229.8
232.1
231.1
240.1
3.8
3.9
4.5
4.3
4.5
3.6
5.4
5.1
4.4
4.0
5.6
4.6
3.9
4.4
5.6
5.9
0.49
0.11
0.29
0.95
0.59
0.07
0.36
1.90
0.49
0.69
0.96
0.76
0.28
0.57
0.18
0.26
0.0114
0.0114
0.0115
0.0116
0.0116
0.0116
0.0117
0.0117
0.0117
0.0118
0.0118
0.0119
0.0119
0.0119
0.0120
0.0121
0.0002
0.0002
0.0002
0.0002
0.0002
0.0002
0.0002
0.0002
0.0002
0.0002
0.0002
0.0003
0.0002
0.0002
0.0002
0.0002
229.1
229.8
230.9
232.3
233.7
233.9
234.6
234.8
235.5
236.8
237.4
238.7
238.8
239.8
241.3
242.7
4.2
4.0
4.2
4.4
4.5
3.7
3.9
4.4
4.4
4.2
4.0
5.0
4.0
4.1
5.0
4.0
Young outliers
MY-01.1
358
MY-01.17
375
1929
2258
5.4
6.0
0.11
0.49
0.0302
0.0341
0.0005
0.0005
191.9
216.0
3.3
3.2
0.13
0.53
0.0094
0.0109
0.0002
0.0002
189.0
218.4
4.1
3.8
>2% common 208Pb
MY-01.18
49
MY-01.13
300
MY-01.5
83
761
987
587
15
3.3
7.1
5.17
1.26
3.78
0.0361
0.0418
0.0353
0.0009
0.0047
0.0008
228.8
263.7
223.8
5.8
28.9
4.9
2.25
2.78
3.30
0.0115
0.0117
0.0117
0.0002
0.0004
0.0002
230.4
235.5
235.2
4.2
7.7
4.7
4f206 [4g208] is the proportion of 206Pb [208Pb] calculated to be common Pb on the basis of measured 204Pb and modeled common Pb
composition (Stacey and Kramers, 1975) at the approximate sample age. All listed Pb isotope data are corrected for common Pb, based on
measured 204Pb/206Pb. Listed uncertainties are 1r and include all components of statistical precision but not the uncertainties in the mean of
the data for the primary reference material 44069 of Aleinikoff et al. (2006) or uncertainties in matrix corrections.
and U concentrations (89–800 ppm and 112–507 ppm,
respectively), with a Th/U ratio of 0.6–1.6. Owing to subsolidus re-equilibration and Pb loss, these analyses yielded
a discordant U–Pb age of 766 ± 25 Ma (MSWD = 2), and
only four samples gave a concordant age of 749 ± 11 Ma
(MSWD = 1.6).
4.2. Trace element composition
The SQ syenites have highly variable trace element
abundances, particularly for Ba, Th, U, Nb, Ta, Pb Zr,
Hf and REE. They are characterized by negative Ta, Sr,
Hf, Sc and Cr anomalies in the primitive mantle-normalized
trace element diagram (Fig. 4a). Positive Ba and Nb anomalies are observed in some samples. The chondrite-normalised REE patterns exhibit LREE enrichment with (La/
Yb)CN ratios ranging from 19 to 56, and negligible to small
positive Eu anomalies (Eu/Eu* = 0.8–1.8; Fig. 4b). Trace
element composition of the SQ carbonatites, reported by
Xu et al. (2010a), shows enrichment in Sr and LREE, and
negative Zr–Hf, Pb, Sc and Cr anomalies in Fig. 4a. The
chondrite-normalised REE patterns of the carbonatites
(Fig. 4b) show a steep negative slope and negligible Ce
and Eu anomalies. Compared to the carbonatites, the syenites contain lower level of Sr, and greater enrichment in
Rb and Ba. Both rock types have similar Sc, V, Ni, Cr
and Ga contents.
4.3. C–O–Sr–Nd–Pb isotopic compositions
The carbon and oxygen isotopic compositions of calcite
separates from the SQ carbonatites show significant variation, ranging from 3.54& to 6.96& d13 CV-PDB and
from 8.62& to 14.36& d18OV-SMOW, respectively (Fig. 5).
Most of these data plot outside of the field of primary,
unaltered carbonatites identified by Taylor et al. (1967).
The d18OV-SMOW values of all SQ samples are higher than
the normal range of mantle values, and some samples show
higher d13CV-PDB values in comparison with the typical
mantle range.
The Sr and Nd isotopic results for the SQ carbonatites
and their constituent minerals (calcite, biotite, pyrite, ferrocolumbite, fluorapatite and monazite) show low level of
radiogenic Sr [(87Sr/86Sr)i 0.7036–0.7040] and little variation in eNd values (from +0.6 to 1.1) approaching to the
Chondritic Uniform Reservoir (CHUR in Fig. 6a). The
model age (TCHUR, Sm–Nd) is mainly in the range of
0.19–0.30 Ga (Table 5). All analyzed carbonatite samples
show a wide range of initial Pb isotopic values. Some agecorrected ratios may be associated with large uncertainties
for the high U, Th contents and 238U/204Pb, 232Th/204Pb
ratios (Table 5). Not uncommonly, the applied corrections
yield unreasonably low or high ratios, resulting in initial
Pb isotopic ratios outside of the ‘normal’ mantle range.
The primary carbonatitic calcite, analyzed in situ by
Table 2
LA-ICPMS U–Pb data from zircons in the South Qinling syenites.
Th (ppm)
U (ppm)
Th/U
207
Pb/206Pb
±1r
207
Pb/235U
±1r
206
Pb/238U
±1r
q
207
Pb/206Pb
Age (Ma)
±1r
207
Pb/235U
Age (Ma)
±1r
206
Pb/238U
Age (Ma)
±1r
66.3
23.2
16.6
48.6
35.1
31.4
23.5
49.0
26.4
77.5
60.0
29.4
27.4
44.9
33.5
58.0
35.8
24.8
27.1
51.9
32.3
62.1
93.1
23.5
352
88.9
108
235
181
156
95.8
322
130
485
457
158
141
215
137
369
190
107
125
273
144
345
800
125
376
160
112
289
209
204
153
283
164
448
370
188
178
272
220
381
214
162
175
332
217
390
507
146
0.93
0.56
0.96
0.81
0.86
0.76
0.62
1.14
0.79
1.08
1.23
0.84
0.79
0.79
0.62
0.97
0.89
0.66
0.72
0.82
0.66
0.89
1.58
0.86
0.0645
0.0645
0.0683
0.0646
0.0648
0.0657
0.0661
0.0652
0.0668
0.0682
0.0674
0.0679
0.0657
0.0635
0.0651
0.0668
0.0655
0.0648
0.0668
0.0667
0.0648
0.0640
0.0649
0.0683
0.0009
0.0013
0.0015
0.0009
0.0010
0.0010
0.0010
0.0011
0.0011
0.0008
0.0010
0.0013
0.0013
0.0010
0.0010
0.0010
0.0010
0.0010
0.0012
0.0009
0.0010
0.0009
0.0009
0.0014
1.1309
1.0050
1.0476
1.0935
1.0862
1.0281
1.0621
1.0504
1.0767
1.1122
1.0581
1.0929
1.0705
1.1018
1.0598
1.0412
1.0995
1.0375
1.0855
1.0381
1.0264
1.0203
1.0273
1.0826
0.0182
0.0220
0.0244
0.0148
0.0172
0.0156
0.0153
0.0177
0.0167
0.0126
0.0179
0.0213
0.0212
0.0172
0.0165
0.0157
0.0175
0.0151
0.0192
0.0135
0.0166
0.0138
0.0141
0.0199
0.1260
0.1123
0.1109
0.1220
0.1211
0.1127
0.1159
0.1163
0.1168
0.1177
0.1129
0.1167
0.1180
0.1252
0.1170
0.1125
0.1209
0.1161
0.1181
0.1128
0.1144
0.1150
0.1140
0.1158
0.0012
0.0011
0.0013
0.0008
0.0010
0.0008
0.0008
0.0009
0.0008
0.0008
0.0008
0.0011
0.0010
0.0010
0.0008
0.0010
0.0009
0.0008
0.0011
0.0007
0.0008
0.0006
0.0007
0.0009
0.5877
0.4319
0.5036
0.4877
0.4978
0.4546
0.4733
0.4476
0.4652
0.5734
0.4402
0.4804
0.4402
0.5189
0.4491
0.5798
0.4429
0.4990
0.5331
0.4664
0.4406
0.4182
0.4329
0.4427
767
767
880
761
769
798
809
789
831
876
850
865
794
724
789
833
791
766
833
828
769
743
769
880
31
39
46
29
228
30
25
37
33
25
32
36
44
33
33
169
33
27
41
28
28
25
28
43
768
706
728
750
747
718
735
729
742
759
733
750
739
754
734
725
753
723
746
723
717
714
718
745
9
11
12
7
8
8
8
9
8
6
9
10
10
8
8
8
8
8
9
7
8
7
7
10
765
686
678
742
737
688
707
710
712
717
690
712
719
761
713
687
736
708
720
689
699
702
696
707
7
6
8
5
5
4
5
5
5
4
5
6
6
6
5
6
5
5
6
4
5
4
4
5
C. Xu et al. / Geochimica et Cosmochimica Acta 143 (2014) 189–206
MY@1
MY@2
MY@3
MY@4
MY@5
MY@6
MY@7
MY@8
MY@9
MY@10
MY@11
MY@12
MY@13
MY@14
MY@15
MY@16
MY@17
MY@18
MY@19
MY@20
MY@21
MY@22
MY@23
MY@24
Pb (ppm)
195
196
C. Xu et al. / Geochimica et Cosmochimica Acta 143 (2014) 189–206
a
b
c
Fig. 3. Weighted average plots of 206Pb/238U (black line) and 208Pb/232Th (grey line) ages derived from monazite SHRIMP data (a), and
whole-rock and mineral Pb–Pb isochron age (b) for the South Qinling carbonatites, and LA-ICPMS analytical U–Pb data for zircons in the
South Qinling syenites with cathodoluminescence images of representative zircons (c). In b, Bt = biotite, Mn = monazite, Ct = ferrocolumbite, Ap = fluorapatite, Py = pyrite; the rest of the data are whole-rock samples. The large MSWD is related to small measurement errors
produced by the double spike method stemming from an increase in the error of regression when the isochron does not pass through the error
range for each sample; MSWD would be significantly lower if the errors for each point were magnified. These data point to either an
underestimated U/Pb ratio, or a geological phenomenon disturbing the primary isochron.
LA-ICPMS, contains tens of ppm Pb, whereas its U and Th
abundances are at or below their detection limit (Table 6),
suggesting that a radiogenic contribution to the Pb budget
of this mineral is negligible. Hence, the Pb budget of the
primary calcite is an accurate measure of the composition
of its parental magma, which in turn, reflects the composition of its mantle source(s). The examined calcite samples
show a wide range of 207Pb/206Pb (0.74–0.88) and 208Pb/206Pb
(1.79–2.15) ratios (Table 6), forming a nearly linear array
between the EM1 and HIMU end-members (Fig. 6b) similar to the so-called East African Carbonatite Line (EACL).
However, combined with Sr–Nd–Pb isotopic data, the SQ
carbonatites form an anomalous isotope pattern (Fig. 6c,
d), which does not follow the HIMU–EM1 mixing line.
Note that, whereas the Sr and Nd isotopic values for the
Bulk Silicate Earth reservoir (BSE) can be estimated with
reasonable accuracy for 235 Ma, the ratios for DMM,
HIMU, EM1 and EM2 end-members cannot be constrained with the same degree of confidence because of
the absence of ocean-island volcanics older than
190 Ma. Thus, we cannot exactly evaluate the isotopic
data for the sources in terms of specific mantle endmembers.
The Sr–Nd isotopic compositions of the SQ syenites
have strongly positive eNd (t) value of 9.6–10.3 and low
radiogenic initial 87Sr/86Sr of 0.7011–0.7034. The Sm–Nd
model age of TCHUR does not fit the syenites with quite
low value of 0.03–0.11 Ga.
The Sr–Nd–Pb isotopic compositions of carbonatites
from the LQ areas, reported by Huang et al. (2009) and
Xu et al. (2011), approach, and trend toward less radiogenic
Sr and Nd values than, the EM1 end-member (Fig. 6a).
Class et al. (2009) argued that EM1 and EM2 components
with narrowly defined isotopic compositions may not exist,
and individual island or volcano with “enriched mantle
affinity” may instead form a trend toward its own unique
end-member. The available data for the SQ and LQ carbonatites plot as clearly distinct clusters (Fig. 6a, c and d), rather
than forming a mixing pattern involving the HIMU and
EM1 end-members. In contrast, Cenozoic carbonatites from
northwestern Pakistan and the Chinese Panxi region in the
Indo-Asian collision zone conform to the well-established
C. Xu et al. / Geochimica et Cosmochimica Acta 143 (2014) 189–206
197
Table 3
Trace element compositions (ppm) of the South Qinling syenites (MY-S17–S20) and carbonatites.
Sc
V
Cr
Ni
Ga
Rb
Sr
Y
Zr
Nb
Ta
Ba
La
Ce
Pr
Nd
Sm
Eu
Gd
Tb
Dy
Ho
Er
Tm
Yb
Lu
Hf
Pb
Th
U
MY-S17*
MY-S18
MY-S19
MY-S20
Carbonatites n = 5*
4.67
190
4.39
22.0
25.0
113
702
25.2
65.7
1151
2.95
15169
154
360
29.5
97.2
12.7
3.59
12.1
1.06
4.47
0.84
2.40
0.31
1.94
0.29
1.93
25.2
10.9
2.24
0.67
101
3.29
11.7
28.7
123
648
86.8
390
387
9.80
1550
462
800
70.9
253
37.1
12.2
43.0
4.58
20.3
3.81
9.67
1.17
7.27
0.92
4.41
37.2
52.5
11.7
4.18
142
1.17
4.74
19.1
31.4
520
18.3
79.6
242
1.39
2520
147
306
32.0
111
14.9
3.77
13.5
1.06
3.83
0.75
2.22
0.27
1.77
0.25
1.26
11.0
5.07
1.27
2.68
40.7
0
3.34
15.9
65.6
1120
16.8
84.6
77.1
1.07
18400
42.2
79.7
9.03
36.0
7.57
5.0
9.44
0.70
3.43
0.70
1.86
0.24
1.48
0.20
2.29
11.7
5.11
0.45
2.33
61.0
2.85
9.44
17.9
11.4
5675
113
62.1
240
5.04
859
503
880
67.2
235
35.8
10.6
33.9
4.08
20.8
3.94
10.7
1.38
8.14
1.15
0.68
21.7
21.2
9.23
Data of Rb, Sr, Ba, Zr, Hf, Nb, Ta, Pb, Th, U and REE from
*
samples were reported by Xu et al. (2010a).
Table 4
C–O isotopic compositions of calcites from the South Qinling carbonatites and calcite veinlets (MY-17) in the syenite.
Sample
MY-0 MY-01 MY-1 MY-4 MY-5 MY-6 MY-7 MY-07 MY-8 MY-10 MY-11 MY-12 MY-13 MY-14 MY-15 MY-16 MY-17
d13CPDB
5.81 4.40
6.96 4.76 3.63 5.30 5.62 6.46
3.78 5.68
4.35
3.54
3.79
4.20
4.49
4.29
4.14
(&)
d18OSMOW 11.04 11.60 9.22 12.21 14.36 12.04 10.92 8.62
10.51 12.49 11.87 11.69 12.43 12.23 11.0
11.50 10.94
(&)
HIMU–EM1 and EM1–EM2 trends, respectively, and show
isotopic compositions different with respect to the SQ and
LQ rocks.
5. DISCUSSION
5.1. Are the SQ carbonatites and syenites genetically related?
The SQ carbonate rocks were emplaced intrusively into
the syenites. In common with calcite carbonatites from
other localities (e.g., Hornig-Kjarsgaard, 1998; Xu et al.,
2007; Chakhmouradian et al., 2008, among many others),
the calcite and fluorapatite in the examined carbonate rocks
are enrichment in Sr and REE (Xu et al., 2010a). The Y/Ho
ratio (average 28.6) of the whole rocks resembles the primary mantle value (28.9; McDonough and Sun, 1995).
The rocks contain mantle-derived Sr–Nd isotopes. Thus,
our field observations and geochemical data support the
interpretation that the examined carbonate rocks are bona
fide carbonatites rather than marble rafts metasomatised
by syenite intrusions.
Models for carbonatite genesis have been debated and
essentially consider two options (e.g., Gittins, 1989): (1)
that carbonatites evolved as a secondary magma during differentiation of a mantle-derived silicate parental melt; or (2)
that carbonatites were derived directly from a mantlederived primary carbonate magma. The first option has
dominated because of close spatial relationship between
carbonatites and alkalic silicate rocks. Although the SQ
carbonatites are intimately associated with the syenites,
the radiochronological data obtained in the present work
imply that the two rock types are not derived from a common parental magma. Zircon crystals from the syenites
define a U–Pb age of 760 Ma, and clearly different from
198
Table 5
Sr–Nd–Pb isotopic compositions of carbonatites and their constituent minerals and syenites (MY-S17–S19) from the South Qinling orogen.
Type
Rb (ppm)
Sr (ppm)
87
MY-0
6.96
0.36
6162
9631
0.19
239
0.35
36.5
0.75
4.06
my-7s
my-8s
my-9s
MY-13
WR
Ap
WR
WR
Bt
Ap
WR
Mn
Ct
WR
WR
Bt
Ct
WR
Py
Py
WR
MY-15
WR
MY-16
WR
MY-17
CV
MY-S17
MY-S18
MY-S19
Samples
MY-0s
MY-01
MY-5
MY-5s
MY-7
MY-0
MY-0s
MY-01
MY-5
MY-5s
MY-7
87
(87Sr/86Sr)i
Sm (ppm)
Nd (ppm)
147
143
eNd (t)
TCHUR (Ga)
0.003
<0.001
0.703650 ± 5
0.703586 ± 6
0.70364
0.70359
3088
625
9800
4009
293
205
<0.001
1.090
<0.001
0.026
0.007
0.057
0.703651 ± 21
0.707534 ± 4
0.703610 ± 18
0.704118 ± 9
0.704050 ± 12
0.704235 ± 7
0.70365
0.70389
0.70361
0.70403
0.70403
0.70405
27.1
14.5
4.13
3724
5338
291
0.021
0.008
0.041
0.703773 ± 5
0.703808 ± 9
0.704189 ± 8
0.70370
0.70378
0.70405
72.1
156
1620
n=9
2327
n=3
3410
n=2
2842
n=6
702
648
520
0.006
0.005
<0.001
0.703685 ± 7
0.703780 ± 5
0.70392* ± 23 (n = 2)
0.70366
0.70376
0.70392
453
867
106
910
163
843
256
54410
770
161
288
397
311
203
122
105
278
0.0817
0.1032
0.0981
0.0738
0.0835
0.1026
0.0912
0.0721
0.0919
0.0992
0.0928
0.0961
0.1003
0.0965
0.0729
0.0732
0.0912
0.512422 ± 4
0.512480 ± 4
0.512446 ± 4
0.512394 ± 3
0.512434 ± 5
0.512478 ± 3
0.512452 ± 3
0.512409 ± 3
0.512434 ± 4
0.512462 ± 4
0.512483 ± 4
0.512474 ± 5
0.512464 ± 7
0.512481 ± 4
0.512415 ± 4
0.512418 ± 4
0.512507 ± 6
0.76
0.28
0.79
1.07
0.58
0.30
0.45
0.73
0.83
0.51
0.09
0.18
0.50
0.06
0.63
0.58
0.61
0.29
0.26
0.30
0.30
0.27
0.26
0.27
0.28
0.30
0.28
0.23
0.25
0.28
0.24
0.28
0.27
0.19
<0.001
0.70361* ± 2 (n = 2)
0.70361
35.7
245
0.0881
0.512487 ± 6
0.31
0.21
<0.001
0.70381* ± 28 (n = 2)
0.70381
33.8
224
0.0913
0.512482 ± 5
0.12
0.23
<0.001
0.70373* ± 1 (n = 2)
0.70373
WR
WR
WR
0.16
0.29
0.01
n=9
0.01
n=3
bdl
n=2
0.01
n=6
113
123
31.4
61.2
148
17.2
111
22.5
143
38.5
6485
117
26.4
44.2
63.1
51.6
32.4
14.7
12.7
41.9
0.460
0.549
0.172
0.705963 ± 14
0.708671 ± 14
0.705284 ± 27
0.70107
0.70271
0.70341
12.7
37.1
14.9
97.2
253
111
0.0790
0.0886
0.0812
0.512577 ± 21
0.512614 ± 9
0.512552 ± 20
10.3
10.1
9.6
0.08
0.03
0.11
Type
U (ppm)
Th (ppm)
238
235
232
206
207
208
204
204
204
204
204
204
204
(206Pb/
Pb)i
(207Pb/
204
Pb)i
(208Pb/
204
Pb)i
(207Pb/
206
Pb)i
(208Pb/
206
Pb)i
74.5
172
41.1
4.63
650
229
60.1
431
472
91.7
320
3909
0.54
1.25
0.30
0.03
4.71
1.66
0.44
3.13
3.42
0.67
2.32
28.3
28.0
129
7.43
17.3
294
44.2
12.9
1.46
36.5
44.1
21.5
217
25.034 ± 1
30.632 ± 3
35.418 ± 3
18.118 ± 1
33.367 ± 2
24.074 ± 1
23.813 ± 1
48.960 ± 3
39.396 ± 3
24.984 ± 1
34.603 ± 5
107.32 ± 24
15.906 ± 1
16.167 ± 2
16.385 ± 1
15.527 ± 1
16.317 ± 1
15.869 ± 1
15.854 ± 1
17.138 ± 1
16.662 ± 2
15.917 ± 1
16.418 ± 2
20.130 ± 1
38.381 ± 3
41.540 ± 6
43.110 ± 5
44.352 ± 1
43.223 ± 4
38.247 ± 1
38.491 ± 3
208.403 ± 1
38.642 ± 5
39.403 ± 3
39.111 ± 1
40.382 ± 13
22.270
24.279
>30
17.946
<10
15.596
21.580
>30
21.872
21.574
22.722
<0
15.765
15.844
16.307
15.518
<15.4
15.438
15.740
>16
15.771
15.744
15.815
<15
38.054
40.029
43.023
44.150
39.786
37.731
38.340
>50
38.215
38.887
38.860
37.846
0.708
0.653
1.709
1.649
0.865
2.460
0.990
0.729
2.419
1.777
0.721
0.730
0.70
1.747
1.803
1.710
WR
Ap
WR
WR
Bt
Ap
WR
Mn
Ct
WR
WR
Bt
20.7
29.7
7.31
2.07
37.2
30.9
39.3
144
2,151
31.9
200
1080
Pb (ppm)
19.3
13.4
14.7
30.5
4.65
9.26
44.7
79.5
377
24.4
49.3
40.3
Rb/86Sr
7.52
21.7
1.28
7.46
16.3
5.78
8.18
0.47
161
14.8
13.0
58.0
Sr/86Sr ±s
U/
Pb
U/
Pb
Th/
Pb
Pb/
Pb ± s
Pb/
Pb ± s
Sm/144Nd
Pb/
Pb ± s
Nd/144Nd ±s
C. Xu et al. / Geochimica et Cosmochimica Acta 143 (2014) 189–206
Samples
Initial Sr, Nd and Pb isotopic ratios of the South Qinling carbonatites and syenites are calculated from Rb, Sr, Sm, Nd, U, Th and Pb contents measured by ICPMS, assuming an age of 235 Ma
and 760 Ma, respectively. The eNd (t) values are calculated based on present-day (147Sm/143Nd)CHUR = 0.1967 and (143Nd/144Nd)CHUR = 0.512638. WR, whole rock; Bt, biotite; Mn, monazite; Ct,
ferrocolumbite; Ap, fluorapatite; Py, pyrite; cal, calcite; CV, calcite veinlets in syenites associated with carbonatites. *Sr isotope was analyzed by calcites from carbonatites using Neptune LA-MCICPMS at Australian National University. The italic initial Pb isotopes are out of a ‘normal’ mantle field, so are not shown in Fig. 6.
MY-7s
MY-8s
MY-9s
Ct
WR
Py
Py
960
21.3
62.8
121
160
54.0
12,540
62,080
72.8
4.71
7.17
6.93
515
27.3
0.32
0.12
3.74
0.20
0.00
0.00
40.4
6.23
0.04
0.01
42.759 ± 4
24.340 ± 6
19.324 ± 2
19.227 ± 2
16.835 ± 2
15.880 ± 1
15.597 ± 1
15.590 ± 1
38.718 ± 6
38.764 ± 13
38.135 ± 4
38.048 ± 5
23.641
23.328
19.312
19.222
15.862
15.829
15.596
15.590
38.246
38.691
38.134
38.048
0.671
0.679
0.808
0.811
1.618
1.659
1.975
1.979
C. Xu et al. / Geochimica et Cosmochimica Acta 143 (2014) 189–206
199
the SHRIMP U–Th–Pb age of monazite from the carbonatites (235 Ma). The syenites show lower initial 87Sr/86Sr and
higher eNd (t) values than the carbonatites (Fig. 6a; Table 5).
Note that some carbonatites and their mineral phases have
quite lower 87Rb/86Sr ratios (<0.001; Table 5), and age-correction cannot affect their initial Sr isotopic composition.
Both liquid immiscibility and crystal fractionation have
been invoked to explain derivation of small-volume carbonate melts from a hybrid alkali-rich carbonate–silicate
magma (Koster van Groos and Wyllie, 1963; Verhulst
et al., 2000; Halama et al., 2005, among many others).
Immiscibility experiments in compositionally diverse silicate–carbonate systems have demonstrated that Pb, Nb,
Th, U and most of the REE partition preferentially into
the silicate liquid, whereas Sr, Ba and F into the conjugate
carbonate fraction (Jones et al., 1995; Veksler et al., 1998,
2012). This pattern of element partitioning is inconsistent
with a much higher content of primary LREE-rich fluorapatite, fluorcarbonates and monazite in the SQ carbonatites
relative to the syenites. Fractionation of feldspars and ferromagnesian minerals from a hybrid carbonate–silicate
magma is a potential alternative mechanism of carbonatite
petrogenesis. Clearly, this process should yield an evolved
melt depleted in trace elements compatible in feldspars, biotite and clinopyroxene and enriched in incompatible elements retained in the liquid phase. The published
experimental studies demonstrated that in silicate melts
(including trachytic compositions), Cr, V and Ni are
strongly compatible with respect to clinopyroxene and biotite (Schmidt et al., 1999; Fedele et al., 2009). In feldspars,
Ga is compatible, whereas Rb is incompatible with DGa 3 DRb (e.g., Bédard, 2006; Macdonald et al., 2010).
Hence, a hybrid magma evolving by fractionation of these
minerals would be expected to produce derivative melts
with significantly lower Ga/Rb ratio and Cr, V and Ni contents relative to the early crystallization products (i.e.,
cumulate syenites). At SQ, this is clearly not the case; the
Ga/Rb ratio is higher in the carbonatites (Ga/Rb = 0.7–
8.8) relative to that in the syenites (0.2–0.6). Both rock
types contain similar contents of compatible transition metals of Cr, V and Ni (Fig. 4a). Based on the above evidence,
we conclude that, despite their intimate spatial association,
the SQ carbonatites are not genetically related to the
syenites.
5.2. Mantle sources, crustal recycling and carbonatite
petrogenesis
The extensive radiogenic isotopic data of carbonatites
worldwide indicate that most young carbonatites
(<200 Ma) have Nd, Sr, and Pb isotopic compositions similar to ocean-island basalts (OIB; Zindler and Hart, 1986;
Hart, 1988), involving HIMU, EM1 and FOZO mantle
components. The HIMU component present in the isotopic
compositions of OIB has been interpreted as Proterozoic
(1–2 Ga) oceanic crust recycled through subduction
(Hofmann, 1997). Based on the similarity of isotopic compositions between OIB and carbonatites, derivation of the
latter from melting of carbonated eclogite has been proposed (Nelson et al., 1988). Published experimental data
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C. Xu et al. / Geochimica et Cosmochimica Acta 143 (2014) 189–206
Table 6
In situ Pb isotopic composition of calcites of carbonatites from the South Qinling orogen.
Samples
U (ppm)
Pb (ppm)
Th (ppm)
(207Pb/206Pb)i ± s
(208Pb/206Pb)i ± s
MY-7 (n = 6)
MY-10 (n = 8)
MY-13 (n = 9)
MY-14 (n = 14)
MY-15 (n = 3)
MY-16 (n = 2)
MY-17 (n = 6)
MY-18 (n = 12)
0.03
0.05
0.01
0.01
bdl
bdl
0.12
0.01
12.1
50.3
13.3
9.99
12.8
9.19
43.0
63.9
0.04
0.01
0.02
0.01
bdl
bdl
0.54
0.12
0.743 ± 22
0.772 ± 10
0.861 ± 16
0.881 ± 11
0.886 ± 2
0.750 ± 22
0.833 ± 8
0.837 ± 8
1.808 ± 83
1.908 ± 25
2.144 ± 31
2.164 ± 23
2.152 ± 16
1.792 ± 23
2.148 ± 9
2.097 ± 19
U, Pb and Th compositions were analyzed by LA-ICPMS. Measured Pb isotopic ratios of calcite are assumed to represent their initial ratios
owing to negligible U and Th contents in the host minerals.
a
b
Fig. 4. Primitive-mantle-normalised trace-element (a) and chondrite-normalised REE abundances (b) of the South Qinling syenites and
carbonatites. The average carbonatite is from Xu et al. (2010a). Normalisation values for primitive mantle and chondrite are from
McDonough and Sun (1995).
Fig. 5. Carbon and oxygen isotopic compositions from the
carbonatites in the South Qinling orogen, shown together with
the field of primary, igneous carbonates (Taylor et al., 1967) and
typical mantle composition (Ray et al., 1999).
suggested that such an origin is feasible (Yaxley and Green,
1994; Hammouda, 2003; Dasgupta et al., 2004, 2005;
Yaxley and Brey, 2004). The experimental work of
Dasgupta et al. (2005) shows that compositions of
eclogite-derived carbonate melts span the range of natural
carbonatites from oceanic and continental settings. The isotopic Sr–Nd–Pb compositions of oceanic carbonatites off
the African coast contain a HIMU-derived signature
(Fig. 3); these compositions have been interpreted to record
melting of Proterozoic subducted oceanic crust stored
within the deep mantle (Hoernle et al., 2002). Petrologic
evidence supports the notion that the carbon cycle involving subducted material may extend to lower-mantle depths
(Walter et al., 2011). However, on the basis of experimental
phase equilibria, Hammouda (2003) argued that carbonates
will most likely be removed from subducted oceanic crust
before it reaches a depth of 300 km. Dasgupta et al.
(2004) and Tsuno and Dasgupta (2012) counter-argued that
melting-induced release of carbonate material from the subducted crust will not occur until the slab reaches at least the
base of the upper mantle and thus carbonates are likely to
be recycled into the deep mantle. On the basis of both
radiogenic and stable isotopic data from carbonatites
worldwide, Bell and Simonetti (2010) interpreted carbonatitic melts as being derived from a sublithospheric source
without any appreciable contribution from crustal materials. The latter authors also emphasized that carbonatites
are virtually absent in subduction-related settings.
There is a general consensus that collision between the
South China Block and Qinling occurred in the Triassic
along the Mianlue suture (Meng and Zhang, 1999; Zhang
et al., 2002; Ratschbacher et al., 2003). This collision was
C. Xu et al. / Geochimica et Cosmochimica Acta 143 (2014) 189–206
a
b
c
d
201
Fig. 6. Comparison of the Sr–Nd–Pb isotopic compositions of carbonatites and seynites in the South Qinling (SQ) and Lesser Qinling (LQ)
orogenic belts and Indo-Asian collision zone. The error bars for in situ Pb isotopic analyses of calcite from the SQ carbonatites reflect in-run
precision with one standard deviation. Data for the LQ-H and LQ-S carbonatites are from Huang et al. (2009) and Xu et al. (2011),
respectively. EACL stands for the East African Carbonatite Line of Bell and Tilton (2001); PaCar and PxCar for Cenozoic Pakistani and
Panxi (China) carbonatites in the eastern Indo-Asian collision zone, respectively (Tilton et al., 1998; Xu et al., 2003; Hou et al., 2006). DMM,
HIMU, EM1 and EM2 are the principal mantle end-member components, as defined by Hart (1988).
preceded by the closure of the Mianlue ocean, which is
manifested in the presence of ophiolite, ocean-island and
island-arc volcanic packages along the southern boundary
of the SQ. The belt of ophiolitic and tectonic mélange
extends E–W for about 160 km (Zhang and Meng, 1995).
Siliceous sedimentary rocks of possible deep water origin
and Early Carboniferous radiolarian fauna (e.g., Albaillella
sp.; Latentifistula cf. ruestae; Entactinia variospina) have
also been recognized in the ophiolite (Feng et al., 1996).
The whole-rock 40Ar/39Ar data for basalts and U–Pb zircon
data for diabase from these ophiolites gave an age range
from 345 to 264 Ma (Lai and Qin, 2010), showing that
the Mianlue crust formed from at least the Early Carboniferous (crust older than 345 Ma may have been subducted)
through the Permian. The important radiometric constraint
on the onset of subduction is a U–Pb zircon age of 246 Ma,
obtained by Qin et al. (2008) for andesite. The SQ carbonatites are located close to the Mianlue suture and were
emplaced at 235 Ma, implying that their parental melt
may be somehow linked to subduction of the Late Paleozoic Mianlue crust.
Higher d18OV-SMOW and d13CV-PDB values were determined for the SQ carbonatites relative to the range of values typically observed in mantle-derived rocks, including
fresh primary carbonatites (Fig. 5). Carbon isotopic composition of carbonatites is less susceptible to modification
by fluid infiltration and other subsolidus processes and,
hence, is a more robust petrogenetic indicator in comparison with d18OV-SMOW (Deines, 1989). The high d13CV-PDB
values recorded in the SQ samples can potentially be
accounted for by liquid immiscibility, fractional crystallization, assimilation of sediments, or source contamination
(Ray et al., 1999). Although these carbonatites are spatially
associated with mantle-derived syenites, the above study
shows they do not have genetically related, ruling out
immiscibility as the driving mechanism for the C-isotopic
excursion. Fractional crystallization can be ruled out
because this process would generate a concerted increase
in d13CV-PDB and d18OV-SMOW values (Ray and Ramesh,
2000), which is not observed in our case (Fig. 5). Assimilation of sediments enriched in heavy C is possible, but is
extremely unlikely to contribute significantly to the C–O
isotopic budget because the SQ rocks show a conspicuous
negative Pb anomaly in their primitive-mantle trace-element pattern (Fig. 4a). Thus, we conclude that crustal carbon recycled to the mantle via subduction of the Mianlue
oceanic crust is the most likely source of the heavy-C
enrichment in the SQ carbonatites.
The origin of carbonatites in orogenic system is disputed. Tilton et al. (1998) suggested that Cenozoic carbonatites from northwestern Pakistan in the Indo-Asian
collision zone were derived from a mantle source similar
to that beneath the East African Rift, i.e., involving upwelling of deep sublithospheric material, and argued that no
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C. Xu et al. / Geochimica et Cosmochimica Acta 143 (2014) 189–206
crustal contribution was involved in the generation of carbonatite parental magmas. Hou et al. (2006) proposed a
petrogenetic model for the Himalayan collision-related carbonatites involving recycling of an old oceanic crust with
pelagic or terrigenous component into a deep mantle
source. Combined the Sr–Nd–Pb isotopic compositions,
the effect of recycling subducted Proterozoic (1.5 Ga) oceanic crust with <10% of sediments into the mantle is modeled in Fig. 7. The results of this modeling show that the
isotopic compositions define a HIMU–EM1–EM2 mixing
trend with different percentages of sedimentary material.
Hence, melting of ancient subducted oceanic crust with a
small sedimentary component is an alternative explanation
for the peculiar isotopic composition of continental carbonatites in the Indo-Asian collision zone. However, the model
yields Sr–Nd–Pb isotopic compositions of Proterozoic oceanic crust plotting above the compositions of the SQ and
LQ carbonatites (Fig. 7). The Paleozoic Mianlue oceanic
crust is characterized by high eNd (8–11) but low (EM1-like)
206
Pb/204Pb ratios (16.90–17.25) (Xu et al., 2002), and is
hence different from the Proterozoic crust used in the modeling. The initial isotopic compositions of the Mianlue crust
and subducted sediments calculated at 235 Ma cannot
explain the CHUR-like Nd and HIMU-like Pb isotopic signatures of the SQ carbonatites. Their formation requires an
additional component with less radiogenic Nd and similarly
radiogenic Pb isotopic characteristics with respect to the
HIMU end-member. Because marine carbonates have low
147
Sm/144Nd (0.11) and high 238U/204Pb ratios (22;
Hoernle et al., 2002), the hypothetical mantle component
with a high 206Pb/204Pb but low 143Nd/144Nd ratios is likely
to be derived from the recycling of Proterozoic carbonates.
These carbonates could be subducted with Proterozoic oceanic crust through deep mantle over a period of 1–2 Ga
before resurfacing in mantle plumes (e.g., Hofmann,
1997) or asthenospheric upwellings. Most of the previously
published isotopic studies of carbonatites appear to favor
plumes to explain both the geochemical characteristics
and tectonic setting of these rocks (Dauphas and Marty,
1999; Bell and Tilton, 2001; Tolstikhin et al., 2002; and references therein). The Qinling orogen lacks any field evidence for plume activity, which is in contrast with the
occurrence of voluminous basalts in the Emeishan igneous
province in the western part of the South China Block,
where the main stage of flood magmatism at 251–253 Ma
(Lo et al., 2002) was associated with the arrival of a mantle
plume (Chung and Jahn, 1995). In this work, we prefer a
model where the magma that produced the SQ carbonatites
originated by melting of the subducted Mianlue crust mixed
with ascending mantle material characterized by high Pb
and low Nd isotopic ratios and derived from a deep-seated
Proterozoic carbonate reservoir via asthenospheric
upwelling.
In comparison with the SQ carbonatites, the LQ rocks
have Sr–Nd–Pb isotopic compositions approaching, and
trending toward less radiogenic Nd value than EM1. The
two groups also differ somewhat in their trace-element composition. The LQ carbonatites are characterized by flat to
Fig. 7. Modeled Sr–Nd–Pb isotopic variations for the carbonatite sources in orogenic belts corresponding to a mixture of recycled oceanic
crust with sediments and of mantle-derived material with delaminated lower crust. The f(Nd, Pb) was calculated using the method described
by Bell and Tilton (2001), i.e., f(Nd, Pb) = [(143Nd/144Nd)2 + (206Pb/204Pb)2]1/2 {sin[arctan(143Nd/144Nd/206Pb/204Pb) + 0.000064]}. (I) (old
OC + sed.) recycled 1.5-Ga oceanic crust (OC; after Rehkämper and Hofmann, 1997) and marine sediments (sed.) at various mass ratios up to
10 wt.%; the average global subducted sediment composition is after Plank and Langmuir (1998); (II) (Ml OC + sed.) recycled Mianlue
oceanic crust (MI OC; after Xu et al., 2002) incorporating up to 8 wt.% of marine sediments (sed.; Plank and Langmuir, 1998); (III)
(MM + LC) mixing of depleted mantle-derived material (MM; estimate from the Mianlue oceanic crust; Xu et al., 2002) with delaminated
lower crust (LC; data from Gao et al., 1998a; Jahn et al., 1999; Lu et al., 2006) at various mass ratios up to 25 wt.%. Data for the Cenozoic
Pakistan (PaCar) and Chinese Panxi carbonatites (PxCar) in the eastern Indo-Asian collision zone, and the LQ-H and LQ-S carbonatites in
Lesser Qinling orogen are shown for comparison. See Fig. 6 for references.
C. Xu et al. / Geochimica et Cosmochimica Acta 143 (2014) 189–206
weakly LREE-enriched chondrite-normalized patterns
[(La/Yb)CN = 1.0–5.5; Xu et al., 2007], which is in marked
contrast with published data for most other carbonatites
[(La/Yb)CN > 20; Chakhmouradian et al., 2008). In addition, these rocks contain economic level of Mo, representing the only known molybdenite deposit associated with
carbonatites worldwide (Xu et al., 2010b). Molybdenite is
not found in the SQ carbonatites, which contain economic
level of REE and are characterized by the normal pattern of
LREE enrichment (Fig. 4b). Therefore, the recorded differences in trace-element and isotopic budget suggest that the
SQ and LQ rocks were derived from compositionally distinct mantle sources.
The LQ carbonatites contain mantle-derived d13CV-PDB
( 5.3% to 7%) and slightly variable d18OV-SMOW (7.6%
to 9.5%) values (Xu et al., 2010b). Recycling of subducted
Paleozoic Mianlue oceanic crust incorporating different
proportions of a sedimentary component cannot explain
the isotopic variation of the LQ carbonatites. It is noteworthy in this context that according to the available geophysical and geochemical evidence, lower crust is absent beneath
the Qinling (Gao et al., 1998b). On the basis of Pb isotopic
evidence for the LQ-S carbonatites, Xu et al. (2011) suggested that the observed crustal thinning resulted from
underthrusting of the southern margin of the North China
Block during its Triassic collision with the South
China Block, followed by gravitational collapse of the
Qinling orogen and delamination of its keel into the mantle.
This type of geodynamic environment has been discussed
previously by Arndt and Goldstein (1989), Kay and
Mahlburg-Kay (1993), and Lustrino (2005). Mixing of the
delaminated lower crust with mantle materials at different
proportions could generate a carbonated source in the
sub-LQ mantle showing less radiogenic Nd isotopic characteristics. The oldest basement rocks exposed in the LQ
region are the Mesoarchean Taihua group comprising
upper-amphibolite- to granulite-facies metamorphic assemblages. Lower-crust xenoliths from the central part of the
North China Block are dominated by two-pyroxene and
garnet-two-pyroxene mafic granulites (Fan and Liu,
1996), implying that these rocks were likely the principal
component of the delaminated lower crust. Using the average chemical composition of granulite from the southern
margin of the North China Block (Gao et al., 1998a,b;
Jahn et al., 1999; Lu et al., 2006), we calculated that a maximum contribution of 15% and 20% of delaminated continental crust is sufficient to produce the unusual isotopic
characteristics of the LQ-S and LQ-H carbonatites, respectively (Fig. 7).
5.3. Tectonic implications
The Qinling orogen merges with the Kunlun and Qilian
orogens to the west (Fig. 1) to compose one of the most
prominent tectonic zones in central East Asia. However,
the exact timing of collision between the South and North
China Blocks along the Qinling orogen remains controversial; contrasting lines of evidence suggest that it either terminated in the Late Triassic (Wang et al., 2007; Qin et
al., 2010; Shen et al., 2014) or began in the Late Triassic
203
and lasted into the Jurassic (Ames et al., 1993;
Ratschbacher et al., 2003; Sun et al., 2002; Hacker et al.,
2004). Emplacement of carbonatites in both the SQ and
LQ places constraints on the regional tectonics following
the closure of the Mianlue ocean. The co-existence of
slightly depleted to enriched mantle sources under the
Qinling belt reflects subduction and recycling of oceanic
and, as we propose in the present work, also lower continental crustal materials into the mantle. Published mineral
Sm–Nd and U–Pb ages indicate that ultrahigh-pressure
metamorphism in the Dabie-Sulu terrane (Fig. 1) took
place between 210 and 245 Ma, peaking at 230–240 Ma
(Li et al., 2000; Ayers et al., 2002; Zheng et al., 2003).
Paleomagnetic studies show that the Permian–Triassic collision between the North and South China Blocks was
accompanied by a 70° clockwise rotation of the latter,
causing the collision to progress from east to west (Zhao
and Coe, 1987). This, in turn, produced extension in the
Dabie-Sulu terrane, leading to rapid exhumation of ultrahigh-pressure metamorphic rocks, and synchronous compression in the Qinling, leading to subduction of the
Mianlue oceanic crust in the Early to Middle Triassic. This
was accompanied by Permian–Triassic plume activity at
Emeishan in the western part of the South China Block
(Chung and Jahn, 1995). The SQ carbonatites were
emplaced in the Late Triassic (235 Ma) at the northwestern margin of the South China Block, and can be convincingly linked to the recycled Mianlue oceanic crust. We
suggest that asthenospheric upwelling triggered decompressional melting of the recycled oceanic crust, producing the
SQ carbonatites. Both asthenospheric activity and melting
could be related to the transition from a transpressional
to transtensional regime at the end of the Late Triassic.
Underthrusting and thickening of the continental crust
beneath the Qinling orogen caused its gravitational collapse
and delamination of the lower crustal material which not
only modified the isotopic budget of the mantle source of
the LQ carbonatites, but also probably caused lithospheric
extension (Jull and Kelemen, 2001). Extensional structures
facilitated melting of the modified mantle source and the
ascent and emplacement of carbonatitic magmas produced
in the process. Therefore, our data conclusively demonstrate that intercontinental collision of the South and North
China Blocks terminated in the Late Triassic.
6. CONCLUSIONS
The carbonatites in the SQ orogenic belt at the northwest margin of the South China Block were emplaced at
233.6 ± 1.7 Ma, i.e., strongly younger than the associated
syenites (766 ± 25 Ma). The initial Sr–Nd isotopic and
trace element compositions do not support that both rock
types have genetically related. The SQ carbonatites were
directly generated by primary carbonate magmas in the
mantle. The rocks exhibit slight depletion in radiogenic
Sr, minor variation in eNd values close to the CHUR, but
a wide range of initial Pb isotopes straddling between the
EM1 and HIMU end-members. Most of their C–O isotopes
plot outside of the field of typical primary igneous carbonates. Emplacement of the SQ carbonatites was preceded by
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C. Xu et al. / Geochimica et Cosmochimica Acta 143 (2014) 189–206
the closure of the Paleozoic Mianlue ocean and collision
between the South China Block and Qinling microplate
along the Mianlue suture. Their isotopic characteristics
are mostly consistent with the melting of subducted oceanic
crust mixed with an ascending deeper mantle component
characterized by high 206Pb/204Pb and low 143Nd/144Nd values and possibly derived from recycled Proterozoic carbonates. In contrast, the Late Triassic (209–221 Ma) LQ
carbonatites at the south margin of the North China Block
contain EM1-like Sr–Nd–Pb and mantle-derived C–O isotopic compositions, distinguishing them from the SQ carbonatites. The isotopic budget of these LQ rocks is best
explained by delamination of lower continental crust
beneath the Qinling following the Late-Triassic collision
between the South and North China Blocks, and incorporation of 15–20% of lower crustal granulites into the subcratonic mantle. The successive emplacement of the SQ and
LQ carbonatites marked a gradual transition from a compressional tectonic regime to an extensional one, indicating
that the collision between the two cratonic blocks terminated in the Late Triassic.
ACKNOWLEDGMENTS
We cordially thank Ian H. Campbell and Charlotte M. Allen
for the laser-ablation ICPMS analyses. We also gratefully acknowledge Prof. Shuguang Li and three anonymous reviewers for their
constructive comments and numerous improvements to this manuscript. Associated Editor Prof. Weidong Sun is thanked for providing useful references. This research was financially supported by the
Chinese National Science Foundation (Nos. 41222022, 41173033);
A.R.C. acknowledges support from the Natural Sciences and Engineering Research Council of Canada.
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