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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: X.Wang3@curtin.edu.au

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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

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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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Pan et al. (2012)

(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/.

ACC

EPTE

D M

ANU

SCR

IPT

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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