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Late Carboniferous N-MORB-type basalts in central Inner Mongolia, China:Products of hydrous melting in an intraplate setting?
Chong-Jin Pang, Xuan-Ce Wang, Bei Xu, Jian-Xin Zhao, Yue-XingFeng, Yan-Yang Wang, Zhi-Wen Luo, Wen Liao
PII: S0024-4937(16)30072-XDOI: doi: 10.1016/j.lithos.2016.05.005Reference: LITHOS 3923
To appear in: LITHOS
Received date: 12 August 2015Revised date: 25 April 2016Accepted date: 9 May 2016
Please cite this article as: Pang, Chong-Jin, Wang, Xuan-Ce, Xu, Bei, Zhao, Jian-Xin,Feng, Yue-Xing, Wang, Yan-Yang, Luo, Zhi-Wen, Liao, Wen, Late Carboniferous N-MORB-type basalts in central Inner Mongolia, China: Products of hydrous melting inan intraplate setting?, LITHOS (2016), doi: 10.1016/j.lithos.2016.05.005
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Late Carboniferous N-MORB-type basalts in central Inner Mongolia, China:
Products of hydrous melting in an intraplate setting?
Chong-Jin Pang1, Xuan-Ce Wang2,3,*, Bei Xu4, Jian-Xin Zhao5, Yue-Xing Feng5,
Yan-Yang Wang4, Zhi-Wen Luo4, Wen Liao4
1 College of Earth Sciences, Guilin University of Technology, Guilin 541004, China
2 The Institute for Geoscience Research (TIGeR), Department of Applied Geology,
Curtin University, GPO Box U1987, Perth, WA 6845, Australia
3 School of Earth science and Resources, Chang’an University, Shannxi, Xi’an
710054, China
4 The Key Laboratory of Orogenic Belts and Crustal Evolution, Ministry of Education,
School of Earth and Space Sciences, Peking University, Beijing 100871, China
5 Radiogenic Isotope Facility, School of Earth Sciences, The University of
Queensland, Brisbane, QLD 4072, Australia
* Corresponding author. Department of Applied Geology, Curtin University, GPO
Box U1987, Perth, WA 6845, Australia. Phone: +61 8 9266 7969 Fax: +61 8 9266 3153 *E-mail: [email protected]
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ABSTRACT
Petrogenesis of the ca. 310 Ma Benbatu basalts in central Inner Mongolia is crucial
for constraining the evolution of the Xing’an Mongolia Orogenic Belt (XMOB),
eastern segment of the Central Asian Orogenic Belt. The Benbatu basalts have low
initial 87Sr/86Sr ratios (0.7042–0.7048), positive εNd(t) (+8.99–+9.24) and εHf(t)
values (+15.38–+15.65), and are characterized by relatively flat rare earth element
patterns and enrichment of Rb, U, Pb, Zr and Hf, but depletion of Nb, Ta, Sr and Ti,
resembling to those of the normal Mid-Ocean-Ridge Basalt (N-MORB). Variations of
trace element ratios (e.g., Sm/Yb and La/Sm) suggest that the basalts were derived
from spinel peridotites, with a melting depth of <60–85 km. The characteristics of
enrichment of Zr and Hf and depletion of Sr distinguish the Benbatu basalts from
typical arc basalts and back arc basin basalts. The arc-like geochemical signatures (i.e.,
enrichment of large ion lithophile elements and depletion of Nb and Ta) are attributed
to hydrated mantle source that may be caused by fluids released from stagnant
oceanic slabs in the mantle transition zone. Integrating geological evidences with
geochemical and isotopic features of the Benbatu basalts, we proposed that these
basalts were produced under an intraplate extensional setting during the Late
Carboniferous. The genesis of the Benbatu basalts therefore argues for the pre-
Carboniferous accretion of the XMOB and highlights the importance of the deep-
Earth recycling water in the generation of the Late Carboniferous magmatism in this
region.
Keywords: Late Carboniferous; the Benbatu basalts; N-MORB-type; Arc-like
geochemical signature; Intraplate setting; Xing’an-Mongolia Orogenic Belt
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1. Introduction
The Central Asian Orogenic Belt (CAOB) extending from Ural area of Russia in the
west, via Mongolia, to Far East area of Russia, and Inner Mongolia and Xing’an areas
of China (Jahn, 2004; Jahn et al., 2000; Windley et al., 2007) is a giant accretionary
orogenic belt between the Siberian craton and the North China and Tarim cratons (Fig.
1a; Hsü et al., 1991; Kröner et al., 2010, 2014; Mossakovsky et al., 1993; Sengör et al.,
1993; Xiao et al., 2003, 2009; Xu et al., 2013, 2015). The eastern segment of the
CAOB covers the Inner Mongolia, Heilongjiang, Jilin and Liaoning provinces in
northern and northeastern China, and is called Xing’an-Mongolia Orogenic Belt
(XMOB, Ren et al., 1980). Different models, especially with respect to the timing of
the final closure of the Paleo-Asian Ocean and associated accretionary processes,
have been proposed to decipher the tectonic evolution of the CAOB between the
Paleozoic and the Mesozoic (Jian et al., 2010; Sengör et al., 1993; Windley et al.,
2007; Xiao et al., 2003, 2009; Xu et al., 2013, 2015; Zhou and Wilde, 2013). For
instance, some authors suggested that the final closure of the Paleo-Asian Ocean and
following amalgamation of CAOB occurred during the Late Permian (Chen et al.,
2000, 2009; Jian et al., 2008, 2010; Xiao et al., 2003, 2009), whereas others proposed
that the southeast part of the Paleo-Asian Ocean was closed by the Middle Devonian
(Xu et al., 2013, 2015). Thus, the Carboniferous sequences in this area are keys to
understanding dynamic transition of the CAOB from the Middle Devonian to Late
Permian, and to testing these different tectonic models.
The primary geochemical and isotopic compositions of mafic rock are served as rock
probe to detect thermochemical state of the deep Earth’s mantle (e.g., Herzberg et al.,
2007; Putirka, 2005; Wang et al., 2007, 2012, 2013, 2015), which are crucial to
understand the evolution of the CAOB. Late Carboniferous volcanic-sedimentary
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rocks are widespread in the XMOB, eastern segment of the CAOB (Fig. 1b), where
mafic rocks are well exposed in the Sonid Youqi and west Ujimqin area (IMBGMR,
1991; Li, 1996; Liu et al., 2013; Pan et al., 2012; Tang et al., 2011). However, the
petrogenesis of these mafic rocks is highly controversial (Liu et al., 2013; Pan et al.,
2012; Tang et al., 2011). For instance, based on their geochemical and isotopic
compositions, Tang et al. (2011) suggested that the Late Carboniferous bimodal
volcanic rocks in the Sonid Youqi were formed in a post-collision extensional setting.
In contrast, others proposed that these volcanic rocks were produced in a subduction
environment only according to their arc-like trace element geochemical signature (i.e.,
depletion of Nb, Ta and Ti; Liu et al., 2013; Pan et al., 2012). This is mainly due to
the lack of systematic geochemical, petrological and mineralogical analyses of these
mafic rocks to constrain their petrogenesis, mantle source characteristics and thus
related geodynamic processes.
Thus, whole rock geochemical, Sr-Nd-Hf isotopic and mineral composition of the
Late Carboniferous mafic rocks in central Inner Mongolia were analyzed in this paper
1) to reveal mantle source characteristics and melting processes of Late Carboniferous
volcanic rocks in central Inner Mongolia; 2) to discriminate tectonic setting for the
generation of these mafic rocks; and 3) to constrain the temporal evolution of the
XMOB.
2. Geological background
The XMOB, bounded by the North China Craton (NCC) to the south, and by the
Siberian Carton to the north (Fig. 1a), consists of the Xing’an-Airgin Sum Block
(XAB), North Orogenic Belt (NOB), Hunshandake Block (HB) and South Orogenic
Belt (SOB) from the north to the south in central Inner Mongolia (Xu et al., 2013,
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2015). The SOB comprises from the north to the south a 40 km-wide E-W trending
fold belt, a 453–445 Ma ophiolitic mélange belt with intensive ductile deformation
and blueschist facies metamorphism, a ca. 600 km-wide 500–450 Ma arc magmatic
belt, and a retroarc foreland basin belt composed of the Silurian turbiditic deposits of
the Xiniwusu Formation and overlying Upper Silurian-Lower Devonian molasses
succession of the Xibiehe Formation (De Jong et al., 2006; Jian et al., 2008; Xu et al.,
2013; Zhang et al., 2013). The NOB is composed of an Early to Middle Paleozoic
(418–482 Ma) arc-pluton complex, a molasses basin filled with ca. 690 m thick red
continental sedimentary rock, a ENE-WSW trending mélange belt with a ca. 380 Ma
high pressure metamorphic age, and a fold belt characterized by greenschist-facies
metamorphism and strong folding deformation, from the north to the south (Xu et al.,
2013). It has been suggested that the SOB was formed by the southward subduction of
the Paleo-Asian Ocean and subsequent collision between the NCC and HB during
500–440 Ma, whereas the NOB was built due to the northward subduction of the
Paleo-Asian Ocean and subsequent collision between the HB and XAB from 500 to
380 Ma (Xu et al., 2013, 2015).
The study area, tectonically belonging to the western part of the HB, is located in the
Qagan Nur area in the northern Sonid Youqi in southern NOB (Fig. 1; Xu et al., 2013).
The oldest sequence in the study area is the Lower Paleozoic Ondor Sum Group,
comprising sericite quartz schist, actinolite schist, blueschist, meta-basalt and
ferriferous quartzite (IMBGMR, 1991; Xu et al., 2013). The Ondor Sum Group
represents continental margin of the HB during the Early Paleozoic and is a
component of mélange matrix (Xu et al., 2013). The Ondor Sum Group is
unconformably overlain by the Benbatu Formation, which is defined as limestone-
dominated sequences with interbedded siliciclastic and volcanic rocks (Li, 1996; Zhu
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et al., 1965). Fossils, such as Fusulinella sp., Lithostrotionella sp., Amygdalophyllum
sp., Koninckophyllum sp., Triticites sp., and Schwaqerina sp., well constrain a Late
Carboniferous age for the Benbatu Formation (Li, 1996; Zhu et al., 1965). The
Benbatu Formation exposed in the Qagan Nur area is mainly composed of shallow-
marine limestone, sandstone and volcanic rocks, and are subdivided into three parts,
namely purple reddish muddy slate in the lower part (C2b1), shallow-marine carbonate
in the middle part (C2b2), and volcanic rocks and interbedded sandstones in the upper
part (C2b3) (Fig. 1c-d; Zhu, 1965). The Benbatu Formation is conformably underlain
by the Amushan Formation, which mainly comprises carbonates with interbedded
basaltic rocks in the West Ujimqin area (e.g., Liu et al., 2013; Fig. 1b-d). Lower
Permian volcanic-siliciclastic successions, assigned as the Dashizhai Formations, are
also widely distributed in the XMOB (Zhang et al., 2008, 2011; Fig. 1b).
The Benbatu Formation, with a thickness of >1000 m, are mainly composed of pillow
basalt, basaltic andesite and tuff in the lower part, but andesite, andesitic porphyrite
and dacite, with interbedded conglomerate, sandstone and carbonate in the upper part
(Fig. 1d). These Benbatu volcanic rocks were also assigned as the Qagan Nur
volcanic rocks in Chinese geological literatures (e.g., Li, 1996). LA-ICPMS zircon U-
Pb isotopic ages constrain that the andesites in upper part were emplaced at 300.9 ±
1.6 Ma (Pan et al., 2012). In contrast, the age of basalts in lower part was poorly
defined. Nonetheless, the Rb-Sr isochron age (313.9 ± 6 Ma) and SHRIMP zircon U-
Pb age (308.5 ± 2.7 Ma) suggest that the basalts in lower part could be formed at 313–
308 Ma (Tang et al., 2011).
To reveal the petrogenesis of the Benbatu volcanic rocks, fourteen samples (14NM79-
96) of the pillow basalts from the lower part of the volcanic sequence and two
andesitic samples (14NM98-99) in the upper part were collected in the northern Sonid
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Youqi for mineralogical, geochemical and Sr-Nd-Hf isotopic analyses (Figs. 1c-d and
2). Except for those with high loss on ignition (>5%), published data of the Benbatu
basalts, andesites and dacites (Pan et al., 2012; Tang et al., 2011) are compared here
to constrain their origin.
3. Petrography
The pillow basalts are mostly fragmented in a quarry outcrop (Fig. 2a), whereas the
andesitic samples on the pasture are massive and relatively fresh (Fig. 2b). The pillow
basalts are characterised by elongate plagioclase laths surrounding very fine-grained
clinopyroxene, exhibiting typical intergranular texture (Fig. 2c-e). Very fine-grained
olivine phenocrysts are observed in sample 14NM82 (Fig. 2d). Andesites are the
dominant lithology in the Benbatu volcanic rocks, and are generally porphyritic (5–20%
phenocrysts). The phenocryst assemblage in andesitic samples mainly comprises
clinopyroxene and plagioclase (Fig. 2f), along with rare quartz. Clinopyroxene
phenocrysts are present as single crystals, from 0.4 to 1.2 mm in size (Fig. 2f).
Plagioclase phenocrysts are mostly subhedral to euhedral, ranging from 0.2 to ca. 1.5
mm in length, showing polysynthetic twin. Groundmass is dominated by microlitic
clinopyroxene and plagioclase (Fig. 2f).
4. Analytical methods
In order to minimize the effect of post-magma alteration, surfaces of samples were
removed, and the freshest central parts chosen for geochemical analysis were crushed
into small chips (<5 mm), which were hand-picked under a binocular microscope to
remove those with surface alteration features and amygdaloidal vesicles. Prior to
being grounded to ˂200 mesh in an agate mill, the purified rock chips were washed in
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0.2N HCl for ˂ 10 minutes and then cleaned with de-ionized water. Whole rock major
and trace elements, Sr-Nd-Hf isotope, and mineral EPMA analyses were conducted in
this study.
Whole rock major element concentrations were determined using an X-ray
fluorescence spectrometer (XRF) on fused glass disks at the Key Laboratory of
Orogenic Belts and Crustal Evolution, Peking University, China. Prior to major
element analysis, loss on ignition (LOI) was determined by heating sample powders
in a muffle furnace at 900°C for several hours before cooled and reweighed.
Analytical uncertainties for majority of major elements were estimated at smaller than
1% from repeatedly analysed Chinese national standards GSR-1 and GSR-3.
The trace element analyses were performed at the Radiogenic Isotope Facility, the
University of Queensland, Australia on a Thermo XSeriesII ICP-MS following the
protocol of Eggins et al. (1997) with modifications as described in Kamber et al.
(2003) and Wei et al. (2014). Sample powders (~60 mg) were dissolved in a distilled
HF-HNO3 (4:1) mixture in Savillex Teflon beakers at 110ºC for 7 days. After drying
down the solutions, sample residues were re-dissolved with double quartz-distilled
concentrated HNO3 followed by 1:1 HNO3 and dried again. The samples were finally
dissolved in an 8 ml 5% HNO3 stock solution. An internal standard containing 6Li,
61Ni, Rh, In, Re, Bi, and 235U was used to monitor and correct instrumental drift.
USGS standard W2 was used as a calibration standard and cross-checked with BIR-1.
The analytical accuracy for most trace elements are typically better than 5%.
Isotope ratios of Sr, Nd and Hf were obtained on the remaining stock solution left
from the trace element analyses. Sr, Nd and Hf chemical separations were performed
by Sr-spec, Thru-spec and LN-spec columns, respectively, following a modified
procedure described by Deniel and Pin (2001), Míková and Denková (2007) and Pin
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and Zalduegui (1997). The 86Sr/88Sr, 143Nd/144Nd and 176Hf/177Hf ratios were
measured in static mode on a Nu Plasma multi-collector inductively coupled plasma
mass spectrometer (MC-ICPMS) at the Radiogenic Isotope Facility, the University of
Queensland following procedures described by Guo et al. (2014) and Wei et al.
(2014). The measured 86Sr/88Sr, 143Nd/144Nd and 176Hf/177Hf ratios were corrected for
mass fractionation by normalizing to 86Sr/88Sr = 0.1194, 146Nd/144Nd = 0.7219 and
179Hf/177Hf = 0.7325, respectively, using an exponential mass fractionation equation.
Standards were used as a monitor of the detector efficiency drift of the instrument.
The SRM-987 standard yielded a measured average of 0.710248 ± 18 (2σ), very close
to the laboratory’s previously obtained long-term average of 0.710249 ± 28 (2σ)
measured using a VG Sector 54 thermal ionization mass spectrometer (TIMS). During
our sample analysis, the Ames Nd metal standard was used as an Nd isotope drift
monitor, and the 143Nd/144Nd data are presented relative to an Ames Nd metal
143Nd/144Nd value of 0.511966, corresponding to a value of 0.512113 ± 9 (2σ, n = 11)
for international standard JNdi-1 standard. For the measurement of 176Hf/177Hf ratios,
a 10 ppm Hf ICP standard solution from Choice Analytical was used as the instrument
drift monitor. During this analysis, the in-house Hf standard monitor was measured
between every 5 samples and gave an average 176Hf/177Hf value of 0.282145 ± 10 (2σ,
n = 13), which was used to correct instrumental drift on measured sample ratios
relative to JMS-475 Hf standard. The 86Sr/87Sr, 143Nd/144Nd and 176Hf/177Hf values of
USGS reference materials BCR-2 and W2 run with our samples are present in Table 2,
which are consistent with the reported reference values (GeoREM:
http://georem.mpch-mainz.gwdg.de/).
Major element analyses of minerals (clinopyroxene) were carried out using a JEOL
JXA-8230 electron microprobe at the Guangxi Key Laboratory of Hidden Metallic
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Ore Deposits Exploration, Guilin University of Technology. The operating conditions
are: 15 kV accelerating voltages, 20 nA beam current, 5 µm beam diameter (or 3 µm
for small grains), and 10 s peak counting time of all elements. All raw data reduction
was carried out using ZAF method.
5. Results
5.1. Major elements
Most of the studied Benbatu basalts have relatively low LOI values (<4.0 wt%; Table
1), except for sample 14NM88, whose whole rock major element contents are not
presented hereafter. The Benbatu basalt samples have a wide range of SiO2 (48.53–
53.75 wt%; normalized to 100% LOI-free), MgO (4.68–5.06 wt%), CaO (9.24–13.86
wt%) and Fe2O3T (11.34–12.54 wt%), with Mg# from 43 to 49 [Mg# =
100×Mg/(Mg+Fe2+) atomic ratio] (Table 1). They are characterized by relatively high
TiO2 (1.81–1.99 wt%) but low Al2O3 (12.83–14.18 wt%) (Fig. 4; Table 1). Samples
show a medium-K to low-K affinity with K2O contents ranging from 0.11 to 0.72 wt%
(Fig. 4; Table 1). Negative correlations between Al2O3, Fe2O3T, CaO and SiO2 are
present, whereas no correlation between TiO2, MgO and SiO2 is shown in the basalts,
excluding 14NM83 with LOI of 4.16 wt% (Fig. 4). In the Zr/TiO2 versus Nb/Y
diagram (Winchester and Floyd, 1976), all Benbatu basalt samples were plotted in the
subalkaline field (Fig. 3b).
In contrast, the Benbatu andesite samples have relatively high SiO2 (60.64–61.89
wt%), low TiO2 (0.74–0.97 wt%), MgO (1.26–1.35 wt%) and LOI (2.33–2.86 wt%)
(Table 1). At a given SiO2, the analyzed samples display relatively lower MgO,
Fe2O3T and TiO2 contents, but higher Al2O3 contents than those of published data (Fig.
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4). Combined with the published data, the Benbatu andesites display negative
correlation between TiO2, MgO, Fe2O3T, CaO and SiO2 (Fig. 4).
5.2. Trace elements
The Benbatu basalts have varied Ni (30.9–41.7 ppm) and Cr concentrations (34.1–
91.7 ppm), much higher than those of andesite samples (1.5–4.6 ppm and 1.3–1.8
ppm, respectively; Table 1). Niobium (Nb) concentrations of the pillow basalts vary
from 1.59 to 1.76 ppm, lower than those of the andesites (Nb = 1.98–4.35 ppm)
(Table 1). Scandium (Sc) concentrations of the basalts range from 36.0 to 39.3 ppm,
much higher than those of the andesites (Sc = 7.78–9.72 ppm; Table 1). The Benbatu
basalts in this study show no correlation between Ni, Cr, Sc, Nb, Rb, Ba, Th, La
concentrations and SiO2 contents (not shown).
The Benbatu basalts exhibit consistent chondrite-normalized rare earth element (REE)
patterns, with low concentrations of most moderately incompatible trace elements and
relatively flat REE pattern ((La/Yb)N = 0.79–0.85), similar to those of normal mid-
ocean ridge basalt (N-MORB) (Sun and McDonough, 1989; Fig. 5a). However, these
samples differ significantly from N-MORB in their enrichment of fluid-mobile
elements (e.g., Rb, Th, U and Pb) relative to fluid-immobile elements (e.g., Nb, Ta
and light REE; e.g., Duggen et al., 2004) in primitive mantle-normalized trace
element patterns (Fig. 5b). These basalt samples also slightly enriched in Zr and Hf,
but weakly depleted in Sr and Ti relative to neighbouring elements (Fig. 5b). In
contrast, the Benbatu andesites have relatively flat REE pattern ((La/Yb)N = 0.65–
1.16), with Eu anomalies (Eu/Eu* = 0.68–0.81; Fig. 5c). These samples are enriched
in Th-U and Zr-Hf, but depleted in Rb, Ba and Ti relative to neighbouring elements
(Fig. 5d). They show either positive or negative anomalies of Pb and Sr (Fig, 5d).
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5.3. Sr-Nd-Hf isotopes
The Benbatu basalts have uniform initial 87Sr/86Sr (87Sr/86Sri) ratios (0.7042–0.7048),
εNd(t) (+8.99–+9.24) and εHf(t) values (+15.38–+15.65; Table 2). In εNd(t) versus
87Sr/86Sri ratio diagram (Fig. 6a), all samples are plotted close to the DMM field
(depleted-MORB mantle; Zindler and Hart, 1986), but differ from the Pacific,
southern Indian and Atlantic MORB by relatively higher 87Sr/86Sr ratios at given
εNd(t) values (Figs. 6a; e.g., Hofmann, 1997; Meyzen et al., 2007; Wang et al., 2013).
The Nd and Hf isotopic covariation of the Benbatu basalts show no linear array (Fig.
6b). The Benbatu basalts have εNd(t) values higher than those of the ca. 280 Ma
Dashizhai volcanic rocks (Fig. 6a; Zhang et al., 2008, 2011). The Benbatu andesites
have similar εHf(t) values (+15.16–+15.37), and 87Sr/86Sri ratios (0.7046–0.7048), but
relatively lower εNd(t) values (+7.36–+8.06) compared to those of the basalts (Table
2).
5.4. Mineral compositions
Clinopyroxenes (Cpx) in four basaltic samples (14NM82, 14NM88, 14NM89 and
14NM92) are analyzed by electron microprobe in this study. Chemical compositions
of the analyzed Cpx are presented on Appendix Table A1. Cations were calculated on
a six-oxygen basis following the program of Sturm (2002), and ferric iron was
calculated using the charge balance method. All Cpx of analyzed samples fall in the
diopside to Mg-augite fields in terms of En-Fs-Wo nomenclature, with end-member
compositions of En26.5-48.2Fs9.8-27.3Wo39.0-48.4 (Fig. 7a; Table A1; Morimoto, 1988). All
Cpx grains were plotted within the Ca-Mg-Fe pyroxene (Quad area) field in the Q-J
diagram (Fig. 7b). Their Mg# values range from 50 to 82.7, and correlate negatively
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with FeOT contents of Cpx. SiO2 contents decrease rapidly from 52 to 45.5 wt% with
the decreasing of Mg# from 82.7 to 69, but slightly increase from 45.5 to 48 wt% with
the decreasing of Mg# from 69 to 50 (Fig. 8a). In contrast, the TiO2, Al2O3, CaO and
MnO contents increase with the decreasing of Mg# from 82.7 to 69, but decrease or
remain nearly constant with the decreasing of Mg# from 69 to 50 (Fig. 8b-f).
6. Discussion
6.1. Effects of alteration and crystal fractionation
In order to reveal source characteristics of parental magma, possible effects of
alteration, crustal contamination and crystal fractionation on whole rock geochemical
compositions are firstly examined here. Most measured samples of the Benbatu
volcanic rocks have relatively low LOI (<4.2 wt%; Table 1), except for sample
14NM88 (LOI = 5.2 wt%), although the alteration of plagioclase and clinopyroxene
can be observed (Fig. 2c-f). The lack of correlation between LOI and major elements
of the Benbatu basalts indicates that their major elements were essentially
immobilized (not shown).
Zirconium (Zr) is suggested to be immobile during low-grade metasomatism and
alteration, and can be used as an alteration-independent index of geochemical
variation for basalts (Polat et al., 2002; Wang et al., 2008, 2010). Thus, bivariate plots
of Zr versus selected trace elements can be used to evaluate mobility of such elements
during alteration. For the Benbatu basalts, REE, high field strength elements (HFSE,
e.g., Nb, Ta, Ti and Hf), V, Sc, Th and Y correlate well with Zr (not shown), implying
that these elements were essentially immobile during alteration (Wang et al., 2010).
Consistency of trace elements and isotopic compositions (Figs. 5 and 6) also excludes
significant effect of alteration on trace element and isotope compositions. The lack of
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correlation between fluid-mobile elements (e.g., Rb, Sr, Ba, U and K) and Zr indicates
that these elements may have been mobile during alteration. Overall, the alteration
played minor roles in whole rock chemical compositions of the Benbatu basalts.
All Benbatu basalt samples are characterized by relatively low MgO (4.68–5.13 wt%),
Cr (34.1–91.7 ppm) and Ni (30.9–41.7 ppm) contents, suggesting that they are not in
equilibrium with mantle peridotites and may have experienced fractional
crystallization (Wang et al., 2012, 2014). Combined with published data, the positive
correlations between Ni, Cr and MgO (Appendix Fig. A1) suggest an importance of
fractional crystallization of olivine and clinopyroxene (± orthopyroxene) in producing
geochemical variation of these basalts. Negative correlations between CaO, Fe2O3T,
Al2O3 and SiO2 (excluding sample 14NM83; Fig. 4) indicate that their parental
magma had also undergone fractionation of clinopyroxene, Fe oxide and plagioclase.
The correlation between major elements of Cpx and associated Mg# can provide
another independent constraint on the assemblage of liquidus phases (Fig. 8; Wang et
al., 2012). The FeOT contents of Cpx in the Benbatu basalts increase rapidly at
Mg#Cpx ≥69 but increase gently at Mg#Cpx ≤69 (Fig. 2d). The TiO2 contents of Cpx
decrease from about 3.5 wt% to 2.0 wt% at Mg#Cpx ≤69 (Fig. 2b), likely resulting
from the Ti-Fe oxide fractionation from parental melt (Stone and Niu, 2009; Wang et
al., 2012). This is consistent with increase of SiO2 with decreasing of Mg#Cpx from 69
to 50, because the Ti-Fe oxide crystallization can cause a SiO2 jump in the residual
melts. The CaO contents increase slightly with decreasing Mg#Cpx at Mg#Cpx >71,
whereas they decrease sharply with falling Mg#Cpx at Mg#Cpx <71 (Fig. 2e). This
suggests that Cpx was the dominant crystallized facies at Mg#Cpx <71 (Mg#melt ≈ 44).
The Al2O3 contents increase with decreasing Mg#Cpx at Mg#Cpx >70, but decrease at
Mg#Cpx <70 (Fig. 8c), suggesting that the fractionation of plagioclase began at Mg#Cpx
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≈ 70 (Mg#melt ≈ 46). In summary, the fractionation of olivine, clinopyroxene, Ti-Fe
oxides and plagioclase played an important role in magma evolution.
Together with published data, the Benbatu andesitic rocks display negative
correlations between TiO2, MgO, Fe2O3T, CaO and SiO2 (Fig. 4), indicating that their
parent magmas had undergone the fractionation of Ti-Fe oxides, olivine and
clinopyroxene. The negative anomaly of Eu (Fig. 5c) implies the fractionation of
plagioclase. Therefore, parental magma of andesitic rocks underwent fractionation of
olivine, clinopyroxene, Ti-Fe oxides and plagioclase.
6.2. Crustal contamination
Crustal contamination is inevitable for mantle-derived melts when they ascend
through continental crust (DePaolo, 1981; Spera and Bohrson, 2001) and can be
identified by the correlation between indices of fractionation and chemical/isotopic
data (Wang et al., 2008, 2014). The lack of correlation between 87Sr/86Sri ratios and Sr
concentrations in the Benbatu basalts (Fig. 9a) is inconsistent with crustal
contamination. Furthermore, all basalt samples have nearly constant εNd(t) values,
showing no significant correlation with SiO2, Nb/La and Th/La (Fig. 9b-d), ruling out
significant crustal contamination during magma ascending. We therefore suggest that
crustal contamination is insignificant in generation of these basalts. In contrast,
together with published data, the negative correlation between SiO2, Th/La and εNd(t)
values (Fig. 9b, d) indicate that the parental magma of andesitic rocks may have
undergone crustal contamination. However, more analyses are required to support this
interpretation.
To summarize, the effect of alteration and crustal contamination on the whole rock
compositions of the Benbatu basalts is insignificant, whereas fractional crystallization
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played an important role in magma differentiation. Because they are sensitive to
source region and possible partial melting processes, the highly incompatible trace
element ratios (e.g., La/Sm and Zr/Hf) of the Benbatu basalts can approximately
reflect the characteristics of parental magma and source region after stripping off
contamination (e.g., Wang et al., 2014).
6.3. Mantle source characteristics
6.3.1. Source mineralogy
Niobium (Nb) concentrations can be served as an index of the degree of partial
melting (e.g., Frey et al. 2000; Wang et al. 2012, 2014). The relationship between
highly incompatible trace element ratios and Nb can be used to evaluate the residual
source mineralogy after ruling out the influence of crustal contamination, post-
magmatic alternation and fractionation crystallization processes (Wang et al. 2012).
The positive correlation between Zr/Hf and Nb in the Benbatu basalts indicates a bulk
solid/melt DZr/Hf <1 (Fig. 10a). Experimental results suggest that Zr and Hf are
compatible in eclogitic garnet with DZr/Hf >1 (Pfänder et al., 2007), whereas the
calculated bulk DZr/Hf values are about 0.3–0.4 for peridotitic sources (Wang et al.
2012). This implies that the Benbatu basalt melts were derived predominantly from
peridotite-dominated source. The nearly constant Tb/Yb ratios with increasing Nb
concentrations suggest that bulk DTb/Yb ≈ 1 (Fig. 10b). It is suggested that DTb/Yb ratios
is <1 for garnet/melt partitioning, whereas DTb/Yb ratios for clinopyroxene/melt,
amphibole/melt and phlogopite/melt partitioning are near unity (Adam and Green
2006; Dalpe and Baker 2000; Latourrette et al. 1995; Wang et al. 2012). This
indicates that garnet could not be the dominant phase in the source mineral
assemblage, but clinopyroxene, amphibole and/or phlogopite may probably be present
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in the generation of the Benbatu basalts. Flattening of heavy REE (HREE) of the
Benbatu basalts (Fig. 5a) also rules out a major role of garnet during partial melting
(Hofmann, 2007).
With partial melting from either spinel- or garnet-bearing peridotites, light REE
(LREE, e.g., La) will be enriched in the melts (Peter et al., 2008). On the other hand,
HREE (e.g., Yb) are incompatible in spinel, but are compatible in garnet (McKenzie
and O’Nions, 1991). Thus, garnet-facies-derived melts will generate higher La/Yb
ratios relative to spinel-facies melts (Peter et al., 2008). Moreover, the presence of
garnet in residual mineral assemblage will cause significant enrichment of middle
REE (MREE; e.g., Dy) relative to HREE (e.g., Yb) during partial melting (Peter et al.,
2008) that was not shown here. Rather, the relatively low Dy/Yb (1.72–1.76) and
La/Yb ratios (1.11–1.19), but relatively high Yb (3.85–4.19 ppm) concentrations of
the Benbatu basalts (Table 1) indicate that partial melting could have occurred within
the spinel stability field (Fig. 10c).In contrast with melting of a garnet lherzolite
source, melts derived from a spinel lherzolite mantle source have similar
MREE/HREE (e.g., Sm/Yb) ratios, but decreasing LREE/MREE (e.g., La/Sm) ratios
with increase of melting degree, plotting within or close to a mantle array defined by
DMM and PM compositions (Fig. 10d; Aldanmaz et al., 2000). The variations of
Sm/Yb and La/Sm ratios thus further constrain that the Benbatu basalts were likely
derived from a spinel lherzolite mantle source (Fig. 10d).
6.3.2. Melting depth and degree
Small degrees of partial melting at high pressures produce magmas with more
normative nepheline (Ne), whereas large degrees of melting at lower pressures
produce magmas with normative hypersthene (Hy) and quartz (Q) (DePaolo and
Daley, 2000). In the Benbatu case, six basalt samples are Ne-normative (normative
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Ne = 0.4–19.3%), while seven basalt samples are Hy-normative, including one Q-
normative sample (14NM84; Table 1). This implies that the primary melts of the
Benbatu basalts could probably have originated from varied melting depths. For
instance, melts of the Q-normative sample could be generated at a melting depth of
<60 km, whereas those of the Ne- and Hy-normative samples were from melting
depths >60 km (e.g., DePaolo and Daley, 2000). The LREE-HREE ratios of the
Benbatu basalts indicate that partial melting could have occurred within the spinel
stability field, which is generally present at depth of about 60–85 km (McKenzie and
O’Nions, 1991; Robinson and Wood, 1998). Therefore, partial melting events to
producing the Benbatu basalts mainly constricted within depth of 60–85 km.
Nonetheless, the consistent trace element and Sr-Nd-Hf isotopic compositions of the
Benbatu basalts suggest that they were derived from a relatively homogenous spinel
lherzolite mantle source, with degrees of partial melting are between 5% and 10%
constrained by REE modelling (Fig. 10d).
6.3.3. Depleted mantle metasomatized by slab-derived fluids
The features of low 87Sr/86Sr ratios (0.7042–0.7048), and coupled positive εNd(t)
(+8.99–+9.24) and εHf(t) values (+15.38–+15.65) suggest a long-term depleted
mantle (MORB like) source for the Benbatu basalts (Fig. 6). By contrast, the
incompatible trace element compositions of the Benbatu basalts exhibit typical
enriched characteristics, as evidenced by enrichment in LILE (e.g., Th and U) and
LREE (e.g., La and Ce) relatively to HFSEs (e.g., Nb and Ta; Fig. 5). For example,
the ratios of Th/Yb and Ta/Yb (or Nb/Yb) are insensitive to fractional crystallization
and/or partial melting, and thus record source variations (Aldanmaz et al., 2000). As
in Fig. 11, the Benbatu basalts display similar Ta/Yb ratios, but higher Th/Yb ratios
than depleted mantle-derived melts, suggesting enrichment of Th in their source
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region (e.g., Aldanmaz et al., 2000; Altunkaynak and Gen, 2008; Gonzalez-Menendez
et al., 2013). It has been demonstrated that such a feature was attributed to source
region rather than crustal contamination (see section 6.2). This implies that the Sr-Nd-
Hf isotope and trace elements are decoupled in depicting the source of the Benbatu
basalts.
The time-integrated effect would result in enriched Sr-Nd-Hf isotopic features if the
enriched reservoirs can be long-term isolated from mantle convection. Thus, if the
enrichment of LILE and LREE relative to HFSE also fractionated Sm/Nd, Lu/Hf and
Rb/Sr, the enriched Sr-Nd-Hf isotopic signatures are expected. The enrichment of Th
relative to deplete mantle-derive melts suggests input of some extent sediments. This
can fractionate mantle Sm/Nd from typical depleted mantle and would result in
enriched Nd isotopes. More importantly, recycling of pelagic sediments into
peridotitic mantle source would result in decoupling Nd-Hf isotopes from the mantle
array (Vervoort et al., 1999). However, the Hf and Nd isotopic compositions of the
Benbatu basalts fall on or close to the Nd-Hf mantle array line (Fig. 6b; Vervoort et
al., 1999). Together with typical depleted mantle-like Nd-Hf isotopic feature, the
enriched incompatible trace element feature should be derived from a young enriched
mantle source without time-integrated Nd-Hf isotopes of Lu/Hf and Sm/Nd
fractionation.
The decoupling between radiogenic isotopes and incompatible trace elements can be
also attributed to mantle hydration processes. Due to significant difference in
solubilities, fluid released from subducted oceanic slabs within the deep-Earth can
impart significant fluid-mobile elements, such as LILE, but little HFSE and REE to
mantle source (e.g., Stolz et al., 1996). Thus, mantle hydration processes can
significantly fractionate LILE to HFSE and REE, but have little effect on internal
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fractionation of HFSE and REE, such as Lu/Hf and Sm/Nd. The most prominent
feature of the Benbatu radiogenic isotope is that although the Nd-Hf isotopes of the
Benbatu basalts are identical with typical depleted mantle-derived melts, such as
MORB, but their Sr isotopes are significant higher than those of MORB (Fig. 6a).
This suggests a hydration enriched mantle reservoir because Sm-Nd and Lu-Hf
isotope systems are insensitive to mantle hydration processes, but Rb is fluid-mobile
element and Rb-Sr would be significantly affected by mantle hydration processes.
Such a relationship between Sr and Nd-Hf isotopes in the Benbatu basalts is crucial
for identifying their hydration mantle reservoir. Thus, the enrichment of fluid-mobile
elements (e.g., Th, U and Pb) relative to fluid-immobile elements (e.g., Nb, Ta, Ti and
REE) suggests that hydrous fluids (water and other fluid phases) played an important
role in metasomatism of the mantle source of the Benbatu basalts. Nonetheless, when
and how the slab-derived fluids metasomatism occurred are poorly constrained but are
keys to an understanding of tectonic processes and basalt petrogenesis (section 6.4.1).
6.4. Hydration mantle melting in an intraplate extensional setting
6.4.1. Evidences for an intraplate extension
The Late Carboniferous Benbatu basalts are characterized by arc-like trace element
signatures, i.e., the enrichment of LILE (e.g., Th and U) and LREE (e.g., La and Ce)
relative to HFSE (e.g., Nb and Ta) as mentioned above (Fig. 5). Such arc-like trace
element features can be observed in back-arc basin basalts (BABB) and true arc
basalts (Kelley et al., 2006; Parman et al., 2011; Saunders and Tarney, 1984; Taylor
and Martinez, 2003), but also present in intracontinental basalts (Wang et al., 2015,
2016). The characteristics of the enrichment of LILE (Rb, Ba, Sr, Th and U) and
LREE (La, Ce, Sm and Nd) relative to HFSE (Nb, Ta, Zr, Hf and Ti) are generally
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attributed to deep-Earth fluid cycling, directly related to subduction processes
(Hawkesworth et al., 1991; McCulloch and Gamble, 1991; Münker et al., 2004; Plank,
2005; Sajona et al., 1996; Stolz et al., 1996), or intraplate processes (Wang, et al.,
2015, 2016). Thus, how to discriminate arc-like intracontinental basalts from BABB
and true arc basalts is crucial to determine their tectonic settings.
Based mainly on the enrichment of LILE and LREE (e.g., Th, U, La, Ce) and the
depletion of HFSE (Nb, Ta and Ti; Fig. 5b), Pan et al. (2012) suggested that the
Benbatu basalts and associated andesites were arc volcanic rocks, inferring a
subduction tectonic setting. Note that all of the documented Benbatu basalts and
associated andesites display prominent positive Zr and Hf anomalies (Fig. 5b), which
are absent in true arc rocks but are common in most typical arc-like intracontinental
basalts, such as Karoo continental flood basalts (Wang et al., 2016). Actually, recent
investigations show that most arc (Tonga, Andes, Luzon, etc.) basalts are marked by
prominent negative Zr-Hf and positive Sr anomalies in primitive mantle normalized
trace element pattern, whereas most arc-like continental (Karoo, Central Atlantic,
Siberian, etc.) basalts do not show these features (Wang et al., 2016 and their figures
4–5). Furthermore, Nb-depleted and Nb-enriched basalts can co-exist in an island arc
volcano (Sorbadere et al., 2013). As a result, instead of the negative Nb and Ta
anomalies, the coupled prominent negative Zr and Hf and positive Sr anomalies are
diagnostic proxies to discriminate true arc basalts from arc-like continental basalts
(Wang et al., 2016). Trace element chemistry of BABBs is transitional between
MORB and island arc basalt, and shows the enrichment of LILE (Rb, Sr, Ba, K and
Th) and the depletion of HFSE (Nb, Ta, Zr, Hf and Ti) as well (Saunders and Tarney,
1984; Stolper and Newman, 1994). The pronounced positive Zr-Hf anomalies of the
Benbatu basalts (Fig. 5b) thus distinguish them from BABBs and true arc basalts.
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Further examinations based on data of less-evolved basalts (MgO ≥8 wt%) show that
the Ti-V, Zr-Zr/Y, Zr-Ti and Ti/V-Zr/Sm-Sr/Nd discrimination diagrams can
successfully distinguish arc-like continental basalts from true arc basalts, because the
primary melts of arc-like continental basalts have different Zr, Sr and Ti compositions
from those of true arc basalts (Wang et al., 2016). For example, a majority of the
Emeishan (87%), Deccan (92%) and Columbia River (85%) continental flood basalts
and the basin and range basalts (88%) is plotted in the intraplate basalt field or offset
from the volcanic arc basalt field on the Zr-Zr/Y discrimination diagram (Wang et al.,
2016). Because their primary magmas had undergone little AFC process, tectonic
setting of the Benbatu basalts can be discriminated using these diagrams. As shown in
Fig. 12, the Benbatu basalts and andesites are plotted within an intraplate basalt field
and show geochemical affinity with the Basin and Range basalts.
Previous studies suggested that the 325–310 Ma quartz diorites in Inner Mongolia
region are arc-related calc-alkaline magmatic rocks formed by the subduction of PAO
during the Late Paleozoic (Chen et al., 2000, 2009; Liu et al., 2009). However, the
presence of Nb-Ta anomalies of those diorites cannot be the diagnostic criteria to
argue for an arc-related origin. Furthermore, the lines of following evidence argue
against the subduction origin of the magmatism. The presence of widespread 290–280
Ma bimodal, shoshonitic to high potassium calc-alkaline volcanic rocks from the
Sonid Zuoqi to West Ujimqin regions in central Inner Mongolia (Fig. 1b) strongly
argues against a subduction tectonic setting, indicating a intraplate extensional setting
during the Early Permian (Zhang et al., 2008, 2011). Recently, geochronological and
geochemical studies show that ca. 308–303 Ma mafic to felsic volcanic rocks in the
Sonid Zuoqi region also display an affinity of shoshonitic to high potassium calc-
alkaline magmatic rock series, implying that the Sonid Zuoqi and adjacent regions
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had undergone intraplate extension during the Late Carboniferous (Gu, 2014; Li et al.,
2014). This is confirmed by evidences from sedimentological and basin analyses.
Stratigraphic data of the Benbatu Formation and equivalents overlying the pre-
Carboniferous rocks (e.g., ophiolite and Precambrian basements) on an unconformity
document an intraplate extensional basin in the Sonid Youqi and adjacent regions
during the Late Carboniferous (Zhao et al., 2015). Fossil assemblage, terrigenous
deposits and associated provenances eliminate the presence of a wide ocean in central
Inner Mongolia during Late Paleozoic (Zhao et al., 2015). Collectively, the arc-like
geochemical signatures observed in the Benbatu basalts most likely reflect hydration
mantle melting within an intraplate tectonic setting caused by deep-Earth fluid cycling
(Wang et al., 2016).
6.4.2. Hydrous melting for generation of the Benbatu volcanic rocks
As magmas crystallize at depth, their major and trace element compositions evolve
along lines defined by their phase equilibria, namely liquid lines of descent (LLD).
Because the stabilities of many minerals (e.g., plagioclase) are highly sensitive to H2O
content, their LLD may document the pre-eruptive volatile content of the magma
(Wang et al., 2016 and references therein). Magma water contents using the Al2O3-
LLD method show that the pre-eruptive H2O contents of the Benbatu basaltic magmas
approach 4.41 wt% (Wang et al., 2016). This implies that mantle hydration played a
crucial role in the generation of the Benbatu basalts. Deep-Earth fluid released from
subducted oceanic slabs can be directly related to subduction processes within arc
environment (e.g., DePaolo and Daley, 2000; Gallagher and Hawkesworth, 1992;
Grove et al., 2012; Niu, 2005; Zimmer et al., 2010), but also can occur within
intraplate tectonic setting (Wang et al., 2015, 2016 and references therein). The
geological, petrological, and geochemical evidences have demonstrated that the
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Benbatu basalts were formed in an intraplate extensional setting. Furthermore, the
decoupling between enriched trace element and typical depleted mantle-like Sr-Nd-Hf
isotopes ruled out the possibility of long-term enriched SCLM origin and suggests a
young hydrated enrichment event in this region.
Recent obtained high-pressure data indicate that subducted oceanic slabs can retain
some extent of their H2O concentrations into the mantle transition zone (MTZ, 410–
660 km; Bizimis and Peslier, 2015; Ivanov and Litasov, 2014). Thus, the MTZ can
serve as a gigantic water tank in the deep-Earth (Wang et al., 2016 and references
therein). The penetration of stagnant slabs from the MTZ into the lower mantle would
subsequently trigger large-scale wet upwelling when the slab avalanches and a
positive buoyancy was generated by accumulation of less-dense fluid phase (Wang et
al., 2015). Water and other fluid phases released from stagnant slabs then hydrated
shallow mantle and introduce fluid-mobile elements to the source (Ivanov and Litasov,
2014, Wang et al., 2015, 2016). Such a water cycle model had been proposed to
account for generation of the arc-like intracontinental basalts (Wang et al., 2015,
2016). Early Paleozoic subduction arc systems are suggested to develop in Inner
Mongolia at 500–410 Ma (Xu et al., 2013 and references therein), nearly 190–100 Ma
prior to the emplacement of the Benbatu basalts, which is consistent with the lifespan
of the stagnant slabs in the MTZ (Wang et al., 2015 and references therein). The
penetration of the stagnant slabs and the intraplate extensional setting may
collectively lead to the MTZ wet upwelling and produce the Benbatu basalts.
Together with their arc-like trace element chemistry, the deep-Earth water cycling
model is thus envisaged to explain the fluid metasomatism and hydration of the
peridotite mantle source of the Benbatu basalts and subsequent partial melting during
the Late Carboniferous.
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The Benbatu andesites have identical εHf(t) values (εHf(t) = +15.16–+15.37) but
slightly lower εNd(t) values (εNd(t) = +7.36–+8.06) compared to those of their
basaltic counterparts (+15.38–+15.65 and +8.99–+9.24 respectively), also implying a
depleted (MORB like) mantle source for their petrogenesis. Direct derivation of the
Benbatu andesites from basaltic magmas by crystal fractionation processes alone
however cannot explain the Nd isotopic variations. Instead, the decreased εNd(t)
values with increasing SiO2 contents in the Benbatu andesites indicate that
contamination of crustal or lithospheric mantle materials may also played an
important role in their petrogenesis. Nonetheless, further studies are required to
constrain the relative importance of crustal or lithospheric mantle materials. The
pronounced positive Zr and Hf anomalies in the Benbatu andesites distinguish them
from true arc volcanic rocks, but indicate an intraplate extensional origin as well.
6.5. Tectonic implication
The Benbatu volcanic rocks serve as crucial deep-Earth probe to determine the Late
Carboniferous tectonic setting of the XMOB. Geochemical and isotopic results in this
study, and available geological evidences well constrain that the Benbatu volcanic
rocks were formed in an intraplate extensional setting during the Late Carboniferous.
Therefore, an integrated magmatism and tectonic evolution is postulated as followed.
During the Early Paleozoic (500–410 Ma), the development of northern and southern
subduction systems in the Paleo-Asian Ocean and subsequent collision between NCC,
XAB and HB formed the SOB and NOB in Inner Mongolia (Xu et al., 2013). The
formation of ca. 380 Ma mélange belt and the deposition of molasse basin on NOB
indicate the closure of the Paleo-Asian Ocean and initiation of a Middle Paleozoic
orogenic belt at ca. 380 Ma (Xu et al., 2013).
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The Late Devonian fossil plant-bearing molasse basin evolved into epicontinental
marine basin during the Late Carboniferous, comprising shallow-marine carbonates,
siliciclastic deposits and volcanic rocks (i.e., the Benbatu and Amushan formations;
Xu et al., 2014; Zhao et al., 2015). This is consistent with field investigations that the
Benbatu basalts and andesites are interbedded with carbonates. Tectonically, the
Benbatu basalts were emplaced at the southern margin of the NOB (Fig. 1), showing a
tight relation with the evolution of the NOB. The shallow melting depths (<60–85 km)
of the Benbatu basalt magmas indicate a relatively thin lithospheric mantle beneath
the XMOB, which is agreed with the intraplate extensional tectonics during the Late
Carboniferous, followed by the development of widespread bimodal magmatism and
rifting during the Permian time (Shao et al., 2014; Xu et al., 2014). Such a regional-
scale extension and lithospheric thinning event may be responsible for triggering
deep-Earth wet upwelling, and subsequent rejuvenated hydration and partial melting
of mantle peridotites during the Late Paleozoic.
7. Conclusions
The Benbatu pillow basalts have a wide range of SiO2 (48.53–53.75 wt%), low MgO
contents (4.68–5.13 wt%) and initial 87Sr/86Sr ratios (0.7042–0.7048), and positive
εNd(t) (+8.99–+9.24) and εHf(t) values (+15.38–+15.65), implying they were derived
from a depleted mantle source. These basalt samples are all featured by the N-MORB-
type REE pattern, and enrichment of Rb, U, Pb, Zr and Hf, but depletion of Nb, Ta, Sr
and Ti, which distinguishes them from the typical arc basalts, implying an intra-plate
setting for their genesis. The arc-like geochemical signature (enrichment of LILE and
depletion of Nb and Ta) is attributed to metasomatism by slab-derived fluid phases
inherited from earlier subduction events. Thus, the geochemical and isotopic results
indicate that the Benbatu basalts were derived from a small degree partial melting of a
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depleted (MORB like), hydrated spinel peridotite mantle source, with a melting depth
of <60–85 km.
The relatively lower εNd(t) values (+7.36–+8.06) but higher SiO2 contents (59–60
wt%) of the Benbatu andesites indicate that assimilation and fractionation
crystallization processes may played an important role in their genesis.
Together with available geological data, the geochemical and isotopic compositions
of the Benbatu volcanic rocks suggest that the closure of Paleo-Asian Ocean and
accretion of the Xing’an Mongolia Orogenic Belt had been completed by Middle
Paleozoic, followed by a post-collisional extension during the Late Paleozoic.
Acknowledgements
We are grateful to two anonymous reviewers for their constructive comments and
suggestions that improved the quality of the manuscript. Editorial assistance by Prof.
T. Wang is gratefully acknowledged. We thank En-Tao Liu in the University of
Queensland, Australia and Yi-Zhi Liu in Guilin University of Technology, China for
technical assistance with diverse analyses. This study has been funded by the Natural
Key Basic Research Program of China (2013CB429806), and the Australian Research
Council (ARC) Future Fellowship (FT140100826) for Xuan-Ce Wang. This work is
also supported by the research grant of State Key Laboratory of Isotope Geochemistry,
Guangzhou Institute of Geochemistry, Chinese Academy of Sciences (SKLIG-KF-13-
07; SKLIG-KF-16-01).
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Figure Captions
Figure 1 Geological maps of the study area and stratigraphic sequence of the Benbatu
Formation. (a) Location of the Central Asian Orogenic Belt (CAOB; Xu et al., 2013).
(b) Tectonic map of western and central Inner Mongolia (modified from Xu et al.,
2013), along with the distribution of Late Carboniferous to Early Permian volcanic-
sedimentary rocks (the Benbatu, Amushan, Dashizhai and Elltu formations; modified
from the 1:50 000 geological map of the Solon-HolinGola region compiled by Tianjin
Institute of Geology and Mineral Resource). (c) Geological map of the Qagan Nur
area in northern Sonid Youqi (modified after 1:20 000 map; Zhu, 1965). (d)
Stratigraphy of the Benbatu Formation in the study area (after Zhu, 1965), along with
fossil and geochronological data.
Figure 2 Photographs of the Benbatu volcanic rocks in the field and microscope
images of the analyzed samples. (a) Pillow basalts in the lower part of the Benbatu
Formation. (b) Massive andesitic rocks in the upper part of the Benbatu Formation.
(c-f) Thin section images of samples 14NM79, 14NM82, 14NM89, 14NM98,
respectively. Ol represents for Olivine, Cpx for clinopyroxene, Pl for plagioclase.
Figure 3 Zr/TiO2 versus Nb/Y diagram for rock type classification (after Winchester
and Floyd, 1976). All samples were normalized to 100% by volatile-free. Literature
data are from Pan et al. (2012) and Tang et al. (2011).
Figure 4 Major elements vs. SiO2 diagrams.
Figure 5 Chondrite-normalized REE pattern and Primitive mantle-normalized
incompatible trace element patterns. Normalized values, N-MORB and E-MORB are
from Sun and McDonough (1989).
Figure 6 Sr-Nd-Hf isotopic compositions of the Benbatu volcanic rocks. (a) 87Sr/86Sr
vs. εNd(t) diagram. Sr-Nd isotopic composition of the Dashizhai volcanic rocks are
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from Zhang et al. (2008, 2011). Sr-Nd isotopic composition range of the southern
Indian MORB and Pacific MORB are from Meyzen et al. (2007) and references
therein. Fields of BSE, EM1 and EM2 are after Zindler and Hart (1986). (b) εNd(t) vs.
εHf(t) diagram. Mantle array line are after Vervoort et al. (1999).
Figure 7 compositional variations of clinopyroxenes. (a) The Pyroxene quadrilateral.
(b) Q-J diagram. Abbreviations and compositions of the end-members are after
Morimoto (1998).
Figure 8 Major elements vs. Mg# of clinopyroxenes.
Figure 9 Diagrams for assimilation-fractionation-crystallization (AFC) discrimination.
(a) Sr vs. 87Sr/86Sr initial ratios. (b) εNd(t) vs. SiO2. (c) εNd(t) vs. Nb/La. (d) εNd(t)
vs. Th/La. The basalt samples were circled with dash lines.
Figure 10 Variations in incompatible trace element ratios constrain source minerals.
(a-b) Variations of trace element ratio (Zr/Hf and Tb/Yb) as a function of partial
melting degree (Nb) in the Benbatu volcanic rocks. (c) Variation of (La/Yb)N and
(Dy/Yb)N of the Benbatu volcanic rocks compared with the melting curves of garnet
and spinel lherzolite mantle source. Melting curves using a primitive mantle
composition (Sun and McDonough, 1989) are after Peter et al. (2008). (d) Variation
of La/Sm and Sm/Yb ratios of the Benbatu volcanic rocks compared with the melt
curves of garnet and spinel lherzolite mantle sources calculated by Aldanmaz et al.
(2000).
Figure 11 Th/Yb vs. La/Yb log-log diagram (after Pearce, 1983) for the Benbatu
volcanic rocks.
Figure 12 (a-d) Plots of Ti vs. V, Zr vs. Zr/Y, Zr-Ti, and Zr/Sm-Sr/Nd-Ti/V for
discriminating the tectonic setting of the Benbatu volcanic rocks. Fields of arc
tholeiitic basalts, calc-alkaline basalts, MORB, continental flood basalts (CFBs),
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Oceanic Island basalts (OIBs), Within Plate basalts (WPB), Island Arc basalts (IAB),
and Basin and Range basalts are after Wang et al. (2016).
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Figure 1
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Figure 2
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Figure 3
Rhyolite
Trachyte
Rhyodacite/Dacite
Andesite
TrachyAndesite
Andesite/Basalt
Alkaline
Basalt
SubAlkaline
Basalt
0.001
0.01
0.1
1
10
0.01 0.1 1 10
Zr/
TiO
*0.0
00
12
Nb/Y
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Figure 4
0.0
0.5
1.0
1.5
2.0
2.5
3.0
TiO
(wt%
)2
12
13
14
15
16
17
18
19
20
Al
O(w
t%)
23
0
1
2
3
4
5
6
Mg
O (
wt%
)
4
6
8
10
12
Fe
O(w
t%)
23
T
1
3
5
7
9
11
13
48 53 58 63 68 73
Ca
O (
wt%
)
SiO (wt%)2
0
1
2
3
4
5
48 53 58 63 68
KO
(w
t%)
2
SiO (wt%)2
73
Shoshonitic
High-K
Medium-K
Low-K
14NM83
(a) (b)
(c) (d)
(e) (f)
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Figure 5
0.1
1
10
100
1000
La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu
Ro
ck/C
ho
nd
rite
0.1
1
10
100
1000
La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu
Ro
ck/C
ho
nd
rite
0.1
1
10
100
1000
RbBa
ThU
NbTa
LaCe
PbPr
SrP
NdZr
HfSm
EuTi
GdTb
DyY
HoEr
TmYb
Lu
Ro
ck/P
rim
itiv
e M
an
tle
0.1
1
10
100
1000
RbBa
ThU
NbTa
LaCe
PbPr
SrP
NdZr
HfSm
EuTi
GdTb
DyY
HoEr
TmYb
Lu
Ro
ck/P
rim
itiv
e M
an
tle
(a) (b)
(c) (d)
E-MORBN-MORB
N-MORB E-MORB
Basalt sample
Andesite-Dacite sample
Tang et al. (2011)
Pan et al. (2012)
Tang et al. (2011)
Basalt sample
Andesite-Dacite sample
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Figure 6
(a)
(b)
BSE
EM1
EM2
-8
-6
-4
-2
0
2
4
6
8
10
12
14
16
0.702 0.704 0.706 0.708 0.710
εN
d(t
)
( Sr/ Sr)i87 86
-8
-4
0
4
8
12
16
20
24
28
32
-8 -6 -4 -2 0 2 4 6 8 10 12 14 16
εH
f(t)
εNd(t)
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Zhang et al. (2008, 2011)
Southern IndianMORB
Pacific MORBSouthern Indian
MORB
PacificM
OR
B
Benbatu volcanic rocks
Dashizhai volcanic rocks
mantle arrayHf = 1.33 Nd+3.19ε ε
basalt
basalt
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Figure 7
0.0
0.5
1.0
1.5
2.0
0.0 0.5 1.0 1.5 2.0
J
Q
Others
Quad
Ca-Na
Na
Ca Si O (Wo)2 2 6
Fe Si O (Fs)2 2 6Mg Si O (En)2 2 6
20
50
45
50
45
20
5
50
enstatite
pigeonite
augite
diopside
5
hedenbergite
ferrosilite
(a)
(b)
14NM88
14NM82
14NM89
14NM92
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Figure 8
45
46
47
48
49
50
51
52
53
SiO
(wt%
)2
0.0
0.5
1.0
1.5
2.0
2.5
3.0
3.5
4.0
TiO
(wt%
)2
5
7
9
11
13
15
Fe
O (
wt%
)2
3
4
5
6
7
Al
O(w
t%)
23
17
18
19
20
21
22
23
49 59 69 79 89
Ca
O (
wt%
)
Mg#
0.0
0.1
0.2
0.3
0.4
0.5
49 59 69 79 89
Mn
O (
wt%
)
Mg#
(a) (b)
(c) (d)
(e) (f)
14NM88
14NM82
14NM89
14NM92
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Figure 9
2
4
6
8
10
0.1 0.2 0.3 0.4 0.5
εN
d(t
)
Nb/La
2
4
6
8
10
48 52 56 60 64 68 72
εN
d(t
)
SiO (wt%)2
0.703
0.704
0.705
0.706
50 150 250 350 450 550 650 750
(S
r/S
r)8
78
6
i
Sr (ppm)
(b)
(c)
(a)
AFC
AFC
AFC
FC
2
4
6
8
10
0.0 0.1 0.2 0.3 0.4 0.5 0.6
εN
d(t
)
Th/La
AFC
FC
(d)
FC
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Figure 10
27
29
31
33
35
37
39
41
43
1 3 5 7 9
Zr/
Hf
Nb (ppm)
0.00
0.05
0.10
0.15
0.20
0.25
0.30
0.35
1 3 5 7 9
Tb
/Yb
Nb (ppm)
0.1
1
10
0.1 1 10
Sm
/Yb
La/Sm
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Pan et al. (2012)
R = 0.52
(a) (b)
(c) (d)
basalt
0.0
0.5
1.0
1.5
2.0
2.5
3.0
3.5
0 10 20 30 40 50
(Dy/Y
b) N
(La/Yb)N
sp+gtIherzolite
50:50
sp-Iherzolite
sp<gt
sp>gt
Depletion Enrichment
sp-Iherzolite
gt-Iherzolite 0.1%1%
5%
10%
20%
30%
50%
1%
5%10%
50%
1%
5%
20%0.1%
5%10%20%
1% 1%5%10%
30%
PMNMORBDMMEMORB
0.5%1%
2.5%
10%
50gt:50sp
Garnet lherzolitemelting curve
Spinel lherzolitemelting curve
La/Sm
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Figure 11
N-MORB
Decreasing Degre
e of
Partial M
elting
MantleArra
y
Average Crust
Su
bd
ucti
on
En
rich
me
nt
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0.01
0.1
1
10
100
0.01 0.1 1 10
Th
/Yb
Ta/Yb
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Figure 12
WTB
VABBasin and Range
Arc tholeiite
calc-alkaline basalts
MORB
CFBs
OIBs
IAB
WPB
ArcTrend
Basin and Range
Basin and Range
MORB
Basin and Range
0
100
200
300
400
500
600
0 5 10 15 20 25
V (
pp
m)
Ti/1000
1000
10000
10 100 1000
Ti (p
pm
)
Zr (ppm)
1
10
10 100 1000
Zr/
Y
Zr (ppm)
Sr/Nd Ti/V
Zr/Sm50000
20
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(a) (b)
(c) (d)
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Table 1 Major element and trace element data of the Benbatu volcanic rocks
Samples Standards
14NM79 14NM81 14NM82 14NM83 14NM84 14NM85 14NM87 14NM88 14NM89 14NM90 14NM91 14NM92 14NM95 14NM96 14NM98 14NM99 GSR-3 Ref. W2 Ref. BHVO-2 Ref.
Rock Basalt Basalt Basalt Basalt Basalt Basalt Basalt Basalt Basalt Basalt Basalt Basalt Basalt Basalt andesite dacite
Major elements (wt%, normalized to 100% volatile-free)
SiO2 50.88 50.44 51.53 48.53 52.54 51.62 50.10 51.44 52.20 51.96 52.70 53.75 53.58 52.24 60.64 61.89 44.70 44.26
TiO2 1.98 1.94 1.88 1.81 1.96 1.91 1.99 1.89 1.96 1.89 1.97 1.95 1.93 2.00 0.97 0.74 2.36 2.43
Al2O3 13.91 14.06 13.83 12.83 13.39 14.12 14.18 13.26 13.48 13.69 13.79 13.23 13.25 13.55 16.37 17.98 13.74 13.62
Fe2O3T 12.38 12.31 11.62 11.34 11.85 11.86 12.54 10.88 11.75 11.37 11.36 11.37 11.19 11.99 6.00 4.93 13.42 13.74
MgO 5.04 4.72 4.69 5.05 4.68 5.13 4.85 5.30 4.72 5.02 4.71 4.98 4.86 5.06 1.35 1.26 7.77 7.55
CaO 10.90 11.77 11.67 13.86 11.00 10.70 12.36 11.77 11.26 10.92 10.14 9.37 9.24 10.03 7.12 4.20 8.73 9.03
K2O 0.23 0.11 0.36 0.39 0.37 0.17 0.58 0.23 0.29 0.26 0.72 0.11 0.13 0.20 0.00 0.01 2.32 2.41
Na2O 4.33 4.30 4.10 5.85 3.87 4.13 3.05 4.87 4.02 4.53 4.30 4.90 5.48 4.60 7.17 8.73 3.37 3.01
MnO 0.15 0.18 0.14 0.16 0.15 0.16 0.16 0.16 0.15 0.16 0.13 0.16 0.16 0.15 0.05 0.05 0.17 0.17
P2O5 0.20 0.19 0.18 0.19 0.19 0.19 0.19 0.19 0.19 0.19 0.18 0.17 0.18 0.18 0.32 0.21 0.94 0.96
LOI 3.19 3.02 4.01 4.16 3.28 3.27 3.55 5.19 3.34 4.09 3.01 3.93 2.33 3.25 2.86 2.33 2.24 2.24
Totala 99.91 99.91 99.91 99.91 99.91 99.90 99.90 99.90 99.90 99.91 99.89 99.91 99.91 99.91 99.93 99.93 99.75 99.41
CIPW (%)
Q 0.00 0.00 0.00 0.00 0.57 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 4.22 0.47
Or 1.35 0.64 2.16 2.31 2.23 1.05 3.49 1.39 1.73 1.54 4.30 0.67 0.80 1.17 0.01 0.06
Ab 35.08 32.41 34.43 14.40 33.18 35.40 26.15 32.24 34.42 37.15 36.77 41.95 45.26 39.42 61.07 74.26
An 18.06 18.97 18.45 7.70 18.25 19.69 23.56 13.80 18.09 16.44 16.39 13.94 11.27 15.92 12.54 9.87
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Ne 1.12 2.40 0.39 19.32 0.00 0.00 0.00 5.09 0.00 0.89 0.00 0.00 0.91 0.00 0.00 0.00
Di 29.30 32.29 32.23 50.05 29.60 27.08 30.92 36.20 30.80 30.65 27.53 26.38 28.01 27.51 17.62 8.18
Hy 0.00 0.00 0.00 0.00 9.60 5.19 5.91 0.00 6.88 0.00 4.36 8.56 0.00 2.57 0.75 4.28
Ol 8.38 6.69 6.01 0.06 0.00 5.12 3.21 5.05 1.56 6.98 4.18 2.10 7.42 6.78 0.00 0.00
Mt 2.45 2.44 2.30 2.25 2.35 2.35 2.49 2.15 2.33 2.25 2.25 2.25 2.21 2.38 1.18 0.97
Il 3.80 3.72 3.61 3.47 3.78 3.67 3.84 3.63 3.76 3.64 3.79 3.74 3.70 3.84 1.86 1.42
Ap 0.48 0.44 0.43 0.45 0.45 0.45 0.45 0.45 0.45 0.45 0.43 0.40 0.42 0.43 0.76 0.49
Trace elements (ppm)
Li 15.5 15.5 18.2 17.3 14.5 16.6 17.4
12.1 18.2 12.0 11.1 12.2 14.9 -0.399 -0.423 9.59 9.3 4.59 4.8
Be 0.88 1.02 0.72 0.72 0.76 0.77 0.92
0.84 0.77 0.74 0.74 0.76 0.80 0.62 0.81 0.75 0.71 1.32 1
Sc 37.9 37.2 37.2 37.6 38.3 37.9 39.2
37.7 37.0 36.0 36.3 39.3 39.3 7.78 9.72 36.07 35.9 32.21 32
V 312 304 297 301 303 311 320
306 300 288 292 310 313 43.9 21.9 261.60 268 315.44 317
Cr 87.7 85.8 87.8 88.4 88.9 89.6 91.7
87.3 85.9 84.6 83.9 34.5 34.1 1.82 1.26 92.79 93 264.02 280
Co 39.5 36.6 40.1 40.2 45.4 53.1 41.3
39.8 42.0 39.2 38.4 43.7 42.1 5.47 7.56 44.53 45 45.48 45
Ni 33.7 34.1 35.8 34.1 39.9 41.7 36.7
31.7 39.5 34.8 32.4 30.9 30.4 1.45 4.56 69.99 72 120.00 119
Samples Standards
14NM79 14NM81 14NM82 14NM83 14NM84 14NM85 14NM87 14NM88 14NM89 14NM90 14NM91 14NM92 14NM95 14NM96 14NM98 14NM99 GSR-3 Ref. W2 Ref. BHVO-2 Ref.
Cu 46.2 55.8 56.4 55.3 58.5 56.0 58.1
54.8 55.2 54.0 52.5 49.3 50.1 8.07 17.1 103.00 105 130.62 127
Zn 88.1 90.8 86.1 88.7 86.7 91.9 93.4 83.5 85.9 77.3 87.1 88.0 88.2 12.3 17.0 77.00 77 103.04 103
Ga 13.6 14.6 12.5 13.6 12.1 12.3 16.7
12.6 12.7 10.9 12.3 12.5 12.0 15.6 14.3 17.42 18 18.66 22
Rb 5.54 2.50 8.62 7.08 9.21 2.81 17.34
7.52 6.55 20.53 3.12 2.80 4.59 0.44 0.72 19.80 21 9.10 9.11
Sr 104 102 147 146 132 116 157
134 167 205 78.6 81.3 150 290 164 194.83 196 392.74 396
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Y 39.8 38.6 36.2 37.5 38.4 38.2 38.6
38.6 36.3 34.8 37.4 36.7 36.6 46.5 49.7 20.11 22 24.45 26
Zr 143 141 138 136 144 142 147
143 136 128 131 140 140 199 250 87.87 92 170.29 172
Nb 1.75 1.68 1.65 1.68 1.73 1.71 1.76
1.74 1.66 1.59 1.64 1.75 1.76 1.98 4.35 7.28 7.5 18.20 18.1
Mo 0.24 0.18 0.17 0.15 0.16 0.23 0.20
0.22 0.26 0.18 0.23 0.19 0.22 0.21 0.66 0.43 0.44 2.75 4
Sn 1.37 1.42 1.31 1.33 1.33 1.29 1.38
1.39 1.33 1.26 1.30 1.22 1.29 0.812 2.23 1.95 2 1.68 1.7
Cs 0.11 0.11 0.16 0.14 0.15 0.11 0.18
0.13 0.14 0.24 0.07 0.11 0.14 0.56 0.69 0.89 0.92 0.10 0.1
Ba 13.7 4.75 9.67 10.2 9.42 6.86 12.1
8.30 9.39 12.9 3.75 6.21 14.4 21.4 21.2 169.68 172 130.29 131
La 4.81 4.46 4.31 4.68 4.55 4.45 4.76
4.69 4.44 4.37 4.48 4.56 4.54 4.74 9.69 10.52 10.8 15.15 15.2
Ce 15.4 14.6 14.2 15.1 15.0 14.7 15.5
15.3 14.5 14.0 14.5 15.1 14.9 26.9 35.2 23.22 23.4 37.55 37.5
Pr 2.67 2.56 2.48 2.60 2.61 2.57 2.67
2.65 2.51 2.43 2.53 2.61 2.56 2.79 4.17 3.03 3 5.37 5.35
Nd 14.0 13.6 13.0 13.7 13.8 13.5 14.0
13.9 13.2 12.8 13.2 13.6 13.4 15.0 19.8 12.91 13 24.26 24.5
Sm 4.73 4.54 4.39 4.54 4.63 4.59 4.74
4.64 4.40 4.29 4.43 4.51 4.50 5.04 5.67 3.27 3.3 6.03 6.07
Eu 1.67 1.62 1.55 1.63 1.64 1.61 1.72
1.65 1.58 1.52 1.57 1.64 1.62 1.55 1.39 1.09 1.08 2.04 2.07
Tb 1.21 1.18 1.13 1.15 1.19 1.18 1.21
1.19 1.13 1.10 1.15 1.15 1.14 1.29 1.33 0.62 0.62 0.99 0.92
Gd 6.30 6.11 5.87 6.03 6.21 6.15 6.29
6.22 5.86 5.70 5.97 5.99 5.96 6.73 6.94 3.71 3.66 6.25 6.24
Dy 7.26 7.01 6.73 6.91 7.12 7.03 7.19
7.12 6.79 6.56 6.82 6.87 6.86 7.87 8.20 3.81 3.79 5.29 5.31
Ho 1.57 1.52 1.45 1.49 1.54 1.52 1.56
1.54 1.47 1.41 1.49 1.48 1.48 1.75 1.85 0.80 0.79 1.00 0.98
Er 4.42 4.28 4.07 4.17 4.31 4.26 4.38
4.32 4.12 3.97 4.18 4.16 4.17 5.15 5.59 2.22 2.22 2.52 2.54
Tm 0.66 0.64 0.61 0.62 0.64 0.64 0.66
0.65 0.62 0.60 0.62 0.62 0.62 0.80 0.88 0.33 0.33 0.34 0.33
Yb 4.19 4.01 3.85 3.93 4.05 4.02 4.15
4.10 3.95 3.78 3.94 3.95 3.98 5.25 6.00 2.06 2.05 2.00 2
Lu 0.62 0.59 0.57 0.58 0.60 0.59 0.61
0.60 0.58 0.55 0.58 0.57 0.58 0.81 0.94 0.30 0.31 0.28 0.27
Hf 3.85 3.75 3.66 3.68 3.79 3.80 3.89
3.82 3.69 3.53 3.64 3.73 3.73 5.56 7.07 2.36 2.45 4.13 4.36
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Ta 0.13 0.13 0.12 0.12 0.13 0.13 0.13
0.13 0.12 0.12 0.13 0.13 0.13 0.13 0.28 0.45 0.47 1.16 1.14
Pb 0.901 0.841 0.919 0.827 0.921 1.00 1.01
0.882 1.66 0.827 0.903 0.759 0.768 0.411 1.29 7.53 7.7 1.50 1.6
Th 0.470 0.454 0.439 0.446 0.466 0.458 0.478
0.472 0.450 0.431 0.443 0.417 0.417 0.952 1.74 2.10 2.17 1.18 1.22
U 0.43 0.34 0.43 0.37 0.37 0.42 0.43 0.41 0.69 0.51 0.54 0.45 0.65 0.44 0.79 0.51 0.51 0.42 0.40
a Before normalization and including loss-on-ignition (LOI).
Ref. = Reference values from GeoREM: http://georem.mpch-mainz.gwdg.de/.
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Table 2 Sr-Nd-Hf isotopic compositions of the Benbatu volcanic rocks
14NM81
14NM82
14NM83
14NM87
14NM95
14NM96
14NM98
14NM99
BCR-2
Ref. W2 Ref.
87Rb/86Sr 0.07095
0.16997
0.14012
0.31958
0.09975
0.08870
0.00442
0.01264
87Sr/86Sr 0.704559
0.705112
0.704904
0.705689
0.705253
0.704963
0.704841
0.704612
0.705001
0.705000
0.706974
2σ 7 8 8 8 6 8 10 8 8 30 8
(87Sr/86Sr)i
0.7042
0.7044
0.7043
0.7043
0.7048
0.7046
0.7048
0.7046
147Sm/144
Nd 0.202484
0.203728
0.200847
0.204177
0.201004
0.202538
0.202813
0.172669
143Nd/144
Nd 0.513123
0.513120
0.513118
0.513122
0.513117
0.513110
0.513027
0.513004
0.512632
0.512636
0.512520
2σ 5 5 4 6 6 5 5 6 5 2 6
(143Nd/14
4Nd)i 0.5127
0.5127
0.5127
0.5127
0.5127
0.5127
0.5126
0.5127
εNd(t) 9.24 9.14 9.20 9.16 9.18 8.99 7.36 8.06
176Lu/177
Hf 0.0225
0.0220
0.0218
0.0221
0.0207
0.0188
176Hf/177
Hf 0.283143
0.283144
0.283147
0.283149
0.283131
0.283126
0.282880
0.282878
0.282733
0.282715
±2σ 4 3
3 3 3 4 4 7 5 3
176Hf/177
Hf(t) 0.2830
0.2830
0.2830
0.2830
0.2830
0.2830
εHf(t) 15.38 15.50 15.65 15.65 15.16 15.37
All errors are 2σ internal precision at the last significant digit. The initial isotopic ratios are calculated at 310 Ma, using relevant element concentration measured by ICP-MS. Reference values are from GeoREM: http://georem.mpch-mainz.gwdg.de/.
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Highlights:
The Late Carboniferous Benbatu basalts have N-MORB-type composition
The basalts derived from partial melting of spinel peridotites
The basalts emplaced in an intraplate extensional setting
Arc-like signatures attribute to hydration mantle melting