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The Early Permian Woniusi Flood Basalts from the Baoshan Terrane

The Early Permian Woniusi Flood Basalts from the Baoshan Terrane,

SW China: Petrogenesis and Geodynamic Implications

CAO Jun1, 2, 3, *

, WANG Xuan3, CHEN Jun

4, ZHU Lixin

5, and Ma Shengming

6

1 State Key Laboratory of Nuclear Resources and Environment, East China University of Technology, Nanchang,

330013, Jiangxi, China 2 Shandong Key Laboratory of Depositional Mineralization & Sedimentary Minerals, Shandong University of

Science and Technology, Qingdao, 266590, Shandong, China 3 School of Earth Sciences, East China University of Technology, Nanchang, 330013, Jiangxi, China 4 State Key Laboratory of Isotope Geochemistry, Guangzhou Institute of Geochemistry, Chinese Academy of

Sciences, Guangzhou, 510640, China 5 China Geological Survey, Beijing, 100037, China 6 Institute of Geophysical and Geochemical Exploration,Chinese Academy of Geological Sciences, Langfang,

065000, Hebei, China

Abstract: The Woniusi flood basalts from the Baoshan terrane, SW China, represent a significant eruption of

volcanic rocks which were linked to the Late Paleozoic rifting of the Cimmeria from the northern margin of East

Gondwana. However, the precise mechanism for the formation and propagation of the rifting is still in debate.

Here we report 40Ar/39Ar dating, whole-rock geochemistry, and Sr–Nd–Pb isotopes for the Woniusi basalts from

the Baoshan terrane of SW China, with the aim of assessing if a mantle plume was related to the formation of the

continent Cimmeria. 40Ar/39Ar dating of the Woniusi basalts yielded ages of 279.5±1.4 Ma and 273.9±1.5 Ma,

indicating they were emplaced during the Early Permian. Whole-rock geochemistry shows that these basalts have

subalkaline tholeiitic affinity, low TiO2 (1.2–2.2 wt.%), and fractionated chondrite-normalized LREE and nearly

flat HREE patterns [(La/Yb)N = 2.86–5.77; (Dy/Yb)N = 1.21–1.49] with noticeable negative Nb and Ta anomalies

on the primitive mantle-normalized trace element diagram. The ɛNd(t) values (-4.76–+0.92) and high (206Pb/204Pb)i

(18.40−18.66) along with partial melt modeling indicates that the basalts were likely derived from a

sub-continental lithospheric mantle (SCLM) source metasomatized by subduction-related processes. On the basis

of a similar emplacement age to the Panjal basalts and Qiangtang mafic dykes and flood basalts in the Himalayas,

combined with a tectonic reconstruction of Gondwana in the Early Permian, we propose that the large-scale

eruption of these basalts and dykes was related to an Early Permian mantle plume that possibly initiated the rifting

on the northern margin of East Gondwana.

Keywords: geochemistry; flood basalts; Baoshan; Gondwana; Early Permian

Email: [email protected]

1 Introduction

It is generally accepted that the opening of the Tethys oceans in the Paleozoic was associated with the rifting and separation of continental fragments (microcontinents or terranes) from Gondwana supercontinent (e.g. Sengor, 1987; Metcalfe, 2013). The Late Paleozoic was a key period of the Paleo-Tethys closure and the Meso-Tethys opening, during which the ribbon-like continent Cimmeria may have separated from the northern margin of East Gondwana (e.g. Ueno, 2003; Metcalfe, 2013). During this time a series of intra-plate, basaltic rocks were emplaced within the Tethyan Himalaya domains of Pakistan, India and China and they are considered to be parts of a Large Igneous Provine (LIP) associated with the breakup of Gondwana (Ernst and Buchan, 2001; Zhu et al., 2010; Liao et al., 2015; Shellnutt et al., 2015; Xu et al., 2016; Zhang and Zhang, 2017; Wang et al., 2019). The Panjal Traps, Abor, Nar-Tsum, Bhote Kosi, and Selong volcanic groups and Qiangtang mafic dykes and flood basalts are amongst the many occurrences of the Late Carboniferous to Early Permian basaltic rocks within the Himalaya (Bhat et al., 1981; Garzanti et al., 1999; Chauvet et al., 2008; Zhu et al. 2010; Ali et al. 2012; Zhai et al. 2013; Wang et al. 2014; Shellnutt et al., 2014, 2015; Xu et al., 2016; Zhang and Zhang, 2017; Wang et al., 2019). Recent studies have also documented large volumes of Late Carboniferous to Early Permian flood basalts and mafic dykes in the Jinji (Wang et al., 2004; Ali et al., 2013; Liao et al., 2015) and Shidian (Huang et al., 2012; Ali et al., 2013) areas in the Baoshan terrane, SW China. However, it is still uncertain whether the regionally extensive volcanic rocks belong to a

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single LIP or whether they represent the juxtaposition of broadly contemporaneous but separate LIP events.

The formation of continental LIPs is a very contentious matter as many are identified as being derived from a mantle plume whereas others are not (Courtillot et al., 1999; Anderson, 2005; Zhang and Dong, 2007; Bryan and Ernst, 2008; Shellnutt et al., 2014). The formation of continental LIPs has also significant implications for continental growth, rifting and breakup (Wang et al., 2008; Bryan and Ferrari, 2013). Previous work suggests the extensive Late Carboniferous to Early Permian basaltic rocks throughout the Himalaya are likely the volcanic expressions of passive lithosphere extension on the northern margin of East Gondwana whereas others suggest they could be the products of basaltic magmatism induced by mantle plume activity (Chauvet et al., 2008; Zhu et al., 2010; Zhai et al., 2013; Shellnutt et al., 2014, 2015; Liao et al., 2015; Xu et al., 2016; Zhang and Zhang, 2017; Wang et al., 2019). Consequently, the origin and tectonic implications of this magmatism remain unclear. However, the close association between LIPs and continental fragmentation means that their petrogenetic interpretation is fundamental to our understanding of the breakup of Gondwana and the evolution of the Tethys ocean (Wilson, 1989; Bryan and Ferrari, 2013).

In this study, we present the first 40

Ar/39

Ar dating and new whole rock Sr–Nd–Pb isotope data for the Early Permian Woniusi basalts from the Baoshan terrane, SW China. These data, along with complementary elemental and mineral chemistry, will help to (i) constrain precise erupted ages of the Woniusi basaltic volcanism, (ii) determine the nature of the sources involved in the genesis of this volcanism, and (iii) reconstruct a model of a mantle plume that was related to Late Carboniferous to Early Permian rifting on the northern margin of East Gondwana.

2 Geological Background

The Sibumasu terrane, which includes the Baoshan, Tengchong and South Qiangtang terranes of western China as well as the terranes of eastern Burma, western Thailand, western Malaysia and Sumatra (Fig. 1a), is an important component of the eastern part of the continent Cimmeria (Metcalfe, 2013). It was separated from Gondwana in Early Permian and accreted to Indochina after the Paleo-Tethyan Ocean was closed in the Middle-Late Triassic (Sone and Metcalfe, 2008).

Fig. 1. (a) Generalized map of the present location of the Cimmerian continent showing the present study area and Early Permian basaltic rocks (green dots) in the South Qiangtang, Panjal Traps, Tethyan Himalaya (modifed after Metcalfe 2013 and Liao et al., 2015). (b) Regional geological sketch map of the Baoshan terrane, northern Sibumasu, modified after Yu (2013). Stars represent the locations of the Jinji and Dongshanpo sampling sections. Occurrence of Daxueshan ultramafc-mafc intrusion that hosts the magmatic sulfde deposit is also shown (modified after Wang et al., 2018).

The Baoshan region of the Sibumasu Gondwana-derived microcontinent has commonly been

referred to as the Baoshan terrane in the literature. The Baoshan terrane represents the northern tip of the Sibumasu terrane (Fig. 1a; Burchfiel and Chen, 2012; Metcalfe, 2013). It is separated from the Indochina block by the Changning-Menglian Paleo-Tethyan suture zone to the east and from the Tengchong terrane by the Cenozoic Gaoligongshan shear zone to the west (Fig. 1b; Burchfiel and Chen, 2012; Li et al., 2015). The basement rocks of the Baoshan terrane are characterized by Late


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Neoproterozoic to Cambrian low-grade metamorphosed siliciclastic, carbonate rocks and some volcanic rock intercalations, and are known as the Gongyanghe Group (Yang et al., 2012; Li et al., 2015). The meta-basalt from the Gongyanghe Group yields zircon U–Pb age of ~499 Ma (Yang et al., 2012). The basement is overlain by a thick sedimentary sequence that includes Silurian, Devonian, Lower Carboniferous, Permian, Jurassic, Tertiary and Quarternary carbonates and clastic rocks (BGMRY, 1981; Xing, 2016). Several important phases of igneous activity have been identified in the Baoshan terrane, including Early Paleozoic (502~448 Ma), Late Carbonifeous to Early Permian (301~282 Ma), and Late Cretaceous to Paleocene (85~60 Ma) events (Wang et al., 2001; Ueno, 2003; Chen et al., 2007; Dong et al., 2013; Li et al., 2015; Liao et al., 2015; Wang et al., 2018). Among these, the Late Carbonifeous to Early Permian phase is the most extensive and considered to be a LIP (i.e. Woniusi LIP) probably related to the inception of rifting of the Baoshan terrane from Gondwana (Wopfner, 1996; Xiao et al., 2003; Liao et al., 2015).

The Woniusi flood basalt province covers an area of ~1.2×104 km

2 within the broad

Baoshan-Shidian-Yongde-Zhengkang region in Yunnan province (Fig. 1b). The Woniusi flood basalt province consists mostly of flood basalts with minor volumes of basaltic pyroclastics, mafic dykes and ultramfic−mafic intrusions (e.g. Daxueshan, Liao et al., 2015; Wang et al., 2018). Stratigraphically, the Woniusi flood basalts unconformably lie over the Upper Carboniferous to Lower Permian Dingjiazhai Formation which consists of a lower unit of glacio-marine diamictites followed by pebby mudstones and an upper unit of shales and bioclastic limestones (Fig. 2 and 3a; Jin et al., 2002). Conodont fossils within the Dingjiazhai formation indicate a deposition age between Gzhelian (304–~299 Ma) and middle Artinskian (287 Ma; Ueno, 2003). Overlying the Woniusi flood basalts is the Lower to Middle Permian Bingma Formation (siltstone and mudstone) (Fig. 2 and 3b). The contact between them is unclear owing to poor exposure (Wang et al., 2001). Some basalt exposures in the Jinji section clearly show pillow structures (Fig. 3a) whereas at Caojian these show relict columnar joints (Liao et al., 2015) indicating that the basalt eruptions likely occurred in both subaqueous and subaerial environments (Kong Manfu, 1990). The thickness of the flood basalts is reported to range from ~40 m in the south (i.e., Yongde; BGMRY, 2004) to ~754 m in the north (i.e., Baoshan; Kong Manfu, 1990). The basalt flows include massive basalts, amygdaloidal basalts, plagioclase-phyric basalts and andesitic basalts, and are generally separated by tuffs, sandstones and shales (Xiao et al., 2003; Yu, 2013; Xu et al., 2015).

Fig. 2. (a) Carboniferous-Permian stratigraphic column in the Baoshan terrane, northern Sibumasu [modified after Jin (2002) and Wang Wei et al. (2004)], showing the stratigraphic positions of the Woniusi basalts (Woniusi Formation). (b) and (c) Two stratigraphic columns of the Woniusi basalts for Jinji and Dongshanpo sections, respectively. The locations of the two samples for 40Ar/39Ar age dating are also shown in (b) and (c).


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3 Sample description and petrography of the basaltic lavas

Two different volcanic outcrops have been sampled. One is located at the Jinji Woniusi area (25°9'7.3''N, 99°17'1.6''E; WNS samples), ~12 km northeast of Baoshan City and the other in Pojiao village of the Youwang Town (24°53'13.3''N, 99°7'10.1''E; DSP samples), ~20 km northwest of Shidian County. The locations of two sampled sections (Jinji and Dongshanpo) are shown in Fig.1b and details of these two sections are summarized in Fig. 2. The two basalt sequences are ~750 m and ~400 m, respectively, and are mainly composed of plagioclase-phyric basalt, massive basalt (Fig. 3c) and amygdaloidal basalt (Fig. 3d) with interbedded tuffs, sandstones and shales. The basalts exhibit porphyritic textures and are characterized by a mineral assemblage of clinopyroxene (~30–45%), plagioclase (~50–60%) and Fe–Ti oxides (~5–10%) (Fig. 3e). The phenocryst assemblage mainly comprises clinopyroxene and plagioclase along with rare Fe–Ti oxides (Fig. 3f). Clinopyroxene phenocrysts are present as subhedral crystals, ranging from 0.1 to 1 mm in size (Fig. 3e). Plagioclase phenocrysts are mostly subhedral to euhedral, ranging from 0.2 to 1.5 mm in length, showing polysynthetic twin (Fig. 3f). Groundmass shows intergranular texture and is mainly composed of fine-grained or aphanitic clinopyroxene, plagioclase and minor Fe–Ti oxides. Replacement of plagioclase by albite and sericite is observed in places. Clinopyroxene has been partially replaced by epidote and chlorite. Amygdules filled with calcite, chlorite and chalcedony are also noted and may be up to ~20–30% in some samples.

Fig. 3. (a) Field photos showing the paraconformable contacts between the Woniusi basalts and the underlying Dingjiazhai Formation. (b) Field photos showing the conformable contacts between the Woniusi basalts and the overlying Bingma Formation. (c) Outcrop photo of massive basalt from the Jinji area. (d) Outcrop photo of amygdaloidal basalt from the Jinji area. (e) Subhedral clinopyroxene phenocryst in the Woniusi basalt; the groundmass is composed of fine grained plagioclase laths and clinopyroxene (sample WNS-16; cross-polarized light). (f) The intersertal and ophitic texture showing clinopyroxene grains within the plagioclase skeletons in the


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Woniusi basalt (sample DSP-19; cross-polarized light).

4 Analytical methods 4.1

40Ar/

39Ar geochronology

Two Woniusi basalts from the Dongshanpo and Jinji sections were selected for dating. Small blocks of each of the two selected samples were then sawn, avoiding any crustal xenoliths, and crushed in a jaw crusher. The fragments were then sieved into fractions from which the groundmass could be carefully picked (~40−60 mesh size fractions). The picked groundmass grains were cleaned with acetone followed by further cleaning with Milli-Q water in an ultrasonic bath for 15 min and then dried at 90 °C. Successively, these grains and monitor ZBH-2506 biotite standard were packaged in aluminium foil, sealed in a aluminium tube and irradiated for 54 h in the 49-2 reactor at the Chinese Academy of Atomic Energy Sciences in Beijing, China. Irradiation flux was monitored using the ZBH-2506 biotite with an age of 132.70 ± 0.01 Ma. Sample J-values were calculated by linear interpolation between two bracketing standards and are included in Supplementary Table 1; a standard was included between every 4 samples in the irradiation tube. Blanks were measured either side of each measurement and used to correct each unknown, and neutron-induced interference reactions using the correction factors (

39Ar/

37Ar)Ca = 6.175×10

-4* (

36Ar/

37Ar)Ca = 2.348×10

-4 and (

40Ar/

39Ar)K =

2.323×10-3

. The

40Ar/

39Ar analyses were performed using a GV5400 mass spectrometer at the Key Laboratory

of Tectonics and Petroleum Resources, Ministry of Education, China University of Geosciences (Wuhan). The detailed analytical procedures are similar to those described by Qiu and Wijbrans (2008). Samples were step-heated using a COHERENT-50W CO2 continuing laser system, with a continuous Nd-YAG 1064 nm infrared laser rastered over the sample for 1min to ensure a homogenously distributed temperature. Gases were gettered for 5 to 8 min using two SAES NP10 getters one at 450 °C and one at room temperature, and a liquid nitrogen cold trap, before inlet to the mass spectrometer. Each analysis was background corrected using blank measurements bracketing every two samples. Data were reduced and age calculations completed using ArArCALC v2.2 software for 40

Ar/39

Ar geochronology (Koppers, 2002). 4.2 Whole-rock analyses

The freshest rock slabs were crushed into small fragments before being further cleaned, then ground into powder ~200 mesh in an agate mortar. Major and trace element analyses were carried out at the State Key Laboratory of Isotope Geochemistry, Guangzhou Institute of Geochemistry, Chinese Academy of Sciences (SKLIG-GIG-CAS). Major elements were analyzed by X-ray fluorescence spectrometer (XRF) on fused glass discs using a Rigaku ZSX100e X-ray fluorescence spectrometer, following the analytical procedures described by Goto and Tatsumi (1996). The analytical uncertainties of the XRF analyses are mostly between 1 and 5%. Trace element concentrations were determined by utilization of PerkinElmer Sciex ElAN 6000 ICP-MS, following the techniques described by Liu et al. (1996). The international standards BHVO-1, BHVO-2, GSR-1, AGV-2 and W-2 were chosen for calibrating element concentrations of the analyzed samples. The analytical precision is better than 5% for elements ˃10 ppm, less than 8% for those ˂10 ppm, and about 10% for transition metals.

Isotope ratios of Sr and Nd of whole-rocks were measured using a Thermo Finnigan Neptune multi-collector ICP-MS at SKLIG-GIG-CAS, following analytical procedures described by Li et al. (2006). The mass fractionation corrections for the isotopic ratios are based on

86Sr/

88Sr = 0.1194 and

146Nd/

144Nd = 0.7219, respectively. Standards were used as a monitor of the detector efficiency drift of

the instrument. The reported 87

Sr/86

Sr and 143

Nd/144

Nd ratios were respectively adjusted to the NBS SRM 987 standard

87Sr/

86Sr = 0.71025 and the JNdi-1 standard

143Nd/

144Nd = 0.512115.

For whole-rock Pb isotopic determinations, about 100 mg powder was weighed into a Teflon beaker, spiked and dissolved in concentrated HF at 180 °C for 7 h. Lead was separated and purified using AG1-X8 resin (200–400 mesh) anionic ion-exchange columns with dilute HBr as eluant. Total procedural blanks were <50 pg Pb. Isotopic ratios were measured by a Thermo Finnigan Neptune multi-collector ICP-MS at the same lab as Sr−Nd isotopic analysis. Thallium was doped as an internal standard to correct the mass fractionation of Pb isotopes during analysis. Repeated analyses of NBS SRM 981 yielded average values of

206Pb/

204Pb = 16.9330 ± 4 (2σ, n=8),

207Pb/

204Pb = 15.4859 ± 4, and

208Pb/

204Pb = 36.6797 ± 13. External precisions are estimated to be less than 0.005, 0.005 and 0.0015.

4.3 Mineral analyses

Major element compositions of the clinopyroxene from samples of the Woniusi basalts were obtained by electron microprobe analysis (EMPA) using a JEOL JXA-8100 Superprobe electron microprobe at SKLIG-GIG-CAS. The accelerating voltage was set at 15 kV with a beam current of 20 nA and 1–2 μm beam diameter. The peak and counting duration was 20 s. The data reduction was performed using ZAF correction procedures.


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Supplementary Table 1 40Ar/39Ar dating results of the Woniusi basalts

Incremental Heating 36

Ar(a) 38

Ar(cl) 39

Ar(k) 40

Ar(r) Age ± 2σ 40

Ar(r) 39

Ar(k)

Steps Laser(%) (V) (V) (V) (V) (Ma) (%) (%)

(a) Sample DSP-10 (Groundmass) T1=279.5±1.4 Ma; T2=283.3±1.4 Ma; T3=279.5±1.6 Ma; T4=279.6±1.6 Ma

16WHA0501A-003 3.2 0.509 0.004 40.557 1109.879 282.9 ± 1.9 88.06 1.18

16WHA0501A-004 3.6 0.310 0.039 52.894 1459.895 285.2 ± 1.5 94.09 1.54

16WHA0501A-005 4.0 0.218 0.000 68.402 1903.050 287.3 ± 1.3 96.72 1.99

16WHA0501A-006 4.5 0.138 0.099 76.202 2145.134 290.4 ± 1.5 98.13 2.22

16WHA0501A-007 5.0 0.085 0.084 85.479 2405.603 290.4 ± 1.2 98.96 2.49

16WHA0501A-008 5.5 0.068 0.120 97.691 2742.287 289.7 ± 1.2 99.26 2.85

16WHA0501B-001 6.0 0.065 0.109 107.126 3015.399 290.4 ± 1.3 99.36 3.12

16WHA0501B-002 6.6 0.035 0.059 95.811 2676.127 288.3 ± 1.2 99.61 2.79

16WHA0501B-003 7.2 0.042 0.127 140.940 3936.572 288.3 ± 1.2 99.68 4.11

16WHA0501B-004 8.0 0.088 0.214 220.624 6126.556 286.8 ± 1.1 99.57 6.43

16WHA0501B-005 9.0 0.066 0.186 230.055 6364.702 285.8 ± 1.1 99.69 6.70

16WHA0501B-006 10.0 0.064 0.237 273.852 7561.346 285.3 ± 1.1 99.74 7.98

16WHA0501B-007 12.0 0.131 0.402 412.087 11286.418 283.1 ± 1.1 99.65 12.01

16WHA0501D-001 15.0 0.063 0.276 241.388 6523.299 279.7 ± 1.1 99.71 7.03

16WHA0501D-002 18.0 0.105 0.226 214.149 5783.122 279.5 ± 1.1 99.46 6.24

16WHA0501D-003 23.0 0.050 0.369 249.751 6743.117 279.4 ± 1.1 99.77 7.28

16WHA0501D-004 28.0 0.175 0.677 374.476 10107.623 279.3 ± 1.1 99.48 10.91

16WHA0501D-005 34.0 0.186 0.885 366.253 9894.835 279.6 ± 1.2 99.44 10.67

16WHA0501D-006 40.0 0.093 0.175 84.440 2278.709 279.3 ± 1.3 98.81 2.46

(b) Sample WNS-5 (Groundmass) T1=273.9±1.5 Ma; T2=260.7±1.3 Ma; T3=270.8±8.0 Ma; T4=270.3±8.1 Ma

16WHA0496-002 2.8 3.579 0.099 7.172 282.805 406.6 ± 25.1 21.10 0.58

16WHA0496-003 3.2 1.242 0.188 21.463 654.041 321.9 ± 3.5 64.05 1.73

16WHA0496-004 3.6 1.748 0.248 44.334 1336.916 318.8 ± 3.2 72.13 3.58

16WHA0496-005 4.0 0.855 0.231 42.012 1165.116 295.2 ± 2.9 82.18 3.39

16WHA0496-006 4.5 0.601 0.216 40.555 1091.656 287.2 ± 2.9 86.00 3.28

16WHA0496-007 5.0 0.374 0.247 40.899 1072.696 280.4 ± 2.8 90.64 3.30

16WHA0496-009 5.5 0.270 0.218 44.839 1151.315 274.9 ± 2.5 93.51 3.62

16WHA0496-010 6.0 0.291 0.238 42.381 1081.442 273.3 ± 3.0 92.62 3.42

16WHA0496-011 6.6 0.296 0.230 43.835 1120.999 273.9 ± 2.4 92.76 3.54

16WHA0496-012 7.2 0.281 0.331 42.408 1084.811 273.9 ± 2.5 92.89 3.43

16WHA0496-013 8.0 0.411 0.803 58.752 1504.418 274.2 ± 1.8 92.53 4.75

16WHA0496-014 9.0 0.387 0.636 59.410 1521.505 274.2 ± 1.9 93.00 4.80

16WHA0496-016 10.0 0.326 0.813 64.413 1641.950 273.1 ± 1.8 94.45 5.21

16WHA0496-017 12.0 0.454 0.968 93.546 2043.173 236.4 ± 1.3 93.82 7.56

16WHA0496-018 15.0 0.550 1.207 103.028 2282.338 239.6 ± 1.2 93.34 8.33

16WHA0496-022 18.0 1.102 4.268 213.789 4737.848 239.6 ± 1.1 93.56 17.28

16WHA0496-023 23.0 0.442 1.580 84.736 1913.097 243.9 ± 1.1 93.60 6.85

16WHA0496-024 30.0 0.478 1.496 88.570 2008.798 244.9 ± 1.2 93.42 7.16


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16WHA0496-025 35.0 0.499 1.279 101.359 2310.906 246.1 ± 1.1 93.99 8.19

T1: Weighted plateau age; T2: total fusion age; T3: isochron age; T4: inverse isochron age. All ages are given at 2σ.

5 Results 5.1

40Ar/

39Ar dating results

The 40

Ar/39

Ar dating results are summarized in Supplementary Table 1 and reported to 2σ. The pure groundmass from sample DSP-10, a massive basalt, yielded a six-step plateau of 279.5 ± 1.4 Ma (Fig. 4a) comprising 44.6% of the total gas released. It has a classic recoil pattern with decreasing measured step age versus temperature. An isochron calculated using nineteen steps produced an age of 279.5 ± 1.6 Ma (Fig. 4b), which is consistent with the plateau age and total fusion age (283.3 ± 1.4 Ma).

The groundmass from sample WNS-5, a plagioclase-phyric basalt, yielded a three-step ‘error-plateau’ (273.9 ± 1.5 Ma; Fig. 4c) comprising 28.8% of the total gas released. This sample has initial

40Ar/

36Ar values (344.2) higher than atmosphere, suggesting excess argon problems. The ‘error

plateau’ is close to a plateau age but is sufficiently disturbed that individual step ages are statistically different from the weighted mean age. The corresponding errorchron age is 270.8 ± 8.0 Ma (Fig. 4d), which is in close agreement with the ‘error plateau’ and may be close to the crystallization age, although it is not completely reliable by the usual criteria.

Fig. 4. (a) 40Ar/39Ar plateau age spectra and (b) 40Ar/36Ar versus 39Ar/36Ar correlation of the groundmass for the Woniusi basalt (Sample DSP-10). (c) 40Ar/39Ar plateau age spectra and (d) 40Ar/36Ar versus 39Ar/36Ar correlation of the groundmass for the Woniusi basalt (Sample WNS-5). All ages are given at the 2σ level. On isochron plots, dark grey squares are data used to determine isochrons.

5.2 Major and trace elements

As noted above, the Woniusi basalts are variably altered in thin section, and this is reflected in their LOI values (LOI = 1.1−4.9 wt.%; Supplementary Table 2). For comparison we have normalized the whole-rock raw data to anhydrous compositions by correcting for LOI, although not all LOI is caused by post-magmatic alteration. We use the normalized values for the following plots and discussions.


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As Na2O and K2O might mobilize during alteration, the Zr/TiO2 versus Nb/Y and SiO2 versus Zr/TiO2 diagrams (Winchester and Floyd, 1977) are utilized for rock classification. On these diagrams Supplementary Table 2 Major element (wt.%) and trace element (ppm) for the Woniusi basalts

Sample DSP-2 DSP-3 DSP-5 DSP-8 DSP-9 DSP-10 DSP-11 DSP-12 DSP-14 DSP-15 DSP-16 DSP-17

Major elements (wt%)

SiO2 50.4 50.2 50.5 54.0 53.1 49.9 50.7 49.8 53.8 56.6 51.7 54.7

TiO2 1.19 1.25 1.43 2.19 1.89 1.64 1.65 1.57 1.56 1.29 1.82 1.39

Al2O3 16.5 16.6 18.2 14.9 14.8 16.1 16.2 16.0 16.1 15.7 14.8 15.2

FeOT 10.2 10.8 9.8 12.0 11.2 10.6 10.3 11.3 10.1 10.1 12.3 10.6

MnO 0.18 0.12 0.12 0.11 0.11 0.13 0.13 0.16 0.12 0.14 0.20 0.19

MgO 7.95 7.16 7.71 3.81 4.20 6.80 6.20 7.52 5.54 4.22 5.51 5.09

CaO 11.2 11.4 9.71 6.26 8.32 10.8 11.1 10.2 9.04 4.90 10.4 9.39

Na2O 2.06 2.10 2.12 5.69 5.50 2.81 2.62 2.35 2.53 5.89 2.31 2.29

K2O 0.24 0.21 0.21 0.63 0.43 0.88 0.78 0.80 0.96 0.92 0.75 0.99

P2O5 0.14 0.14 0.16 0.44 0.39 0.29 0.28 0.25 0.24 0.17 0.22 0.14

L.O.I 3.06 2.21 3.89 2.55 2.71 2.10 1.89 2.60 3.55 2.90 1.68 1.24

Total 99.6 99.6 99.7 99.7 99.7 99.6 99.6 99.7 99.7 99.7 99.6 99.6

Trace elements (ppm)

Sc 37.5 37.7 37.8 41.1 37.6 34.2 35.7 36.3 15.1 39.7 43.9 40.0

V 256 264 259 194 194 244 273 270 251 245 324 264

Cr 472 562 567 440 467 335 204 326 96.8 340.8 57.6 59.2

Co 45.4 46.3 49.5 49.3 43.9 42.6 42.6 50.0 34.6 32.9 45.8 41.8

Ni 165 183 195 149 151 134 106 174 30.8 94.0 43.6 35.3

Cu 92.3 107.8 112.9 84.4 104.5 71.9 36.1 22.3 69.0 65.9 188 68.6

Ga 17.9 18.2 20.3 13.0 20.4 18.1 18.9 19.9 18.6 16.7 20.7 19.3

Rb 3.13 3.36 3.27 10.70 6.46 16.4 13.5 18.0 13.3 30.0 11.3 19.6

Sr 151 169 165 120 113 290 285 245 210 220 183 163

Y 24.1 24.7 25.5 28.1 21.5 20.9 24.0 24.7 25.1 30.0 37.2 31.0

Zr 104 108 121 134 100 112 118 128 160 156 175 139

Nb 7.17 7.47 8.57 14.5 10.7 13.4 13.9 14.1 17.7 16.5 15.0 9.30

Ba 117 112 115 86 56 281 242 234 348 225 244 243

La 10.3 10.6 12.7 20.3 14.7 15.9 16.4 16.6 21.1 19.5 20.2 18.0

Ce 23.6 24.4 27.9 47.8 33.4 33.3 34.7 35.0 45.1 41.3 43.6 37.7

Pr 3.03 3.16 3.61 6.14 4.41 4.44 4.70 4.62 5.71 5.05 5.52 4.89

Nd 13.3 13.9 15.5 26.6 19.3 18.8 20.0 19.4 22.7 20.4 23.2 20.0

Sm 3.63 3.74 4.14 6.39 4.76 4.45 4.82 4.70 5.22 4.99 5.91 5.00

Eu 1.27 1.32 1.46 2.34 1.81 1.71 1.74 1.61 1.69 1.44 1.72 1.45

Gd 4.19 4.28 4.65 6.51 4.86 4.62 5.05 4.93 5.42 5.43 6.48 5.35

Tb 0.74 0.75 0.82 1.03 0.78 0.72 0.81 0.82 0.87 0.91 1.12 0.91

Dy 4.69 4.78 5.15 6.05 4.63 4.29 4.91 5.01 5.28 5.76 7.00 5.71

Ho 1.00 1.04 1.10 1.22 0.94 0.88 1.02 1.04 1.11 1.22 1.50 1.22

Er 2.72 2.82 3.03 3.13 2.45 2.26 2.67 2.76 2.97 3.33 4.04 3.32

Tm 0.41 0.42 0.45 0.44 0.35 0.33 0.39 0.40 0.44 0.49 0.60 0.50

Yb 2.58 2.65 2.83 2.72 2.14 2.04 2.44 2.47 2.74 3.06 3.80 3.13

Lu 0.40 0.40 0.43 0.40 0.32 0.31 0.38 0.38 0.42 0.46 0.57 0.48

Hf 2.58 2.70 3.03 3.33 2.45 2.60 2.88 3.09 3.85 3.88 4.44 3.65

Ta 0.46 0.49 0.55 0.96 0.72 0.86 0.90 0.91 1.17 1.21 1.04 0.67

Pb 2.51 2.47 2.40 3.28 2.18 1.88 2.32 2.55 4.57 5.27 5.62 7.18

Th 1.91 1.94 2.17 2.56 1.83 1.83 1.93 2.49 3.67 4.91 5.33 5.55

U 0.25 0.42 0.22 0.62 0.29 0.28 0.40 0.30 1.15 1.06 0.98 0.93

All oxide contents of the samples have been recalculated to 100% on a volatile-free basis with all Fe as FeOT (total FeO).


9

Supplementary Table 2 (continued)

Sample DSP-18 DSP-19 WNS-9 WNS-11 WNS-12 WNS-16 WNS-19 WNS-20 BHVO-1 Ref.


SiO2 54.9 54.0 52.4 50.9 51.8 50.0 51.5 51.0

TiO2 1.39 1.80 1.53 1.86 1.59 1.64 1.68 1.75

Al2O3 15.2 14.6 15.0 16.2 16.7 15.3 16.0 15.8

FeOT 10.6 11.1 13.4 10.6 10.2 10.6 10.0 10.9

MnO 0.24 0.17 0.16 0.13 0.23 0.15 0.12 0.12

MgO 4.43 5.50 5.24 6.57 5.17 8.28 7.04 7.31

CaO 9.18 9.61 9.36 8.05 10.3 11.5 9.65 9.45

Na2O 2.64 2.29 2.41 4.12 2.64 1.94 2.61 2.58

K2O 1.29 0.74 0.32 1.16 1.07 0.35 1.04 0.77

P2O5 0.14 0.22 0.16 0.33 0.24 0.26 0.31 0.29

L.O.I 1.06 2.17 4.88 2.75 1.15 4.42 3.09 3.64

Total 99.6 99.6 99.8 99.7 99.6 99.8 99.7 99.7


Sc 41.1 40.8 41.9 32.9 34.2 36.5 35.8 35.7 31.6 31.4

V 282 293 284 246 249 258 247 254 316 314

Cr 60.3 139 136 223 259 459 369 364 260 288

Co 41.7 41.3 44.0 42.1 36.2 47.9 42.7 41.1 44.9 44.9

Ni 36.9 42.7 52.2 116 63.6 180 123 128 121 120

Cu 17.3 117 71.4 34.6 212 78.5 25.4 14.0 134.3 137.2

Ga 19.6 19.9 18.7 18.0 17.6 17.5 18.2 17.6 21.4 21.3

Rb 45.3 11.5 8.51 28.7 21.3 3.5 21.6 16.7 9.3 9.5

Sr 170 190 166 356 226 289 254 235 395 399

Y 31.4 37.5 27.6 23.1 24.2 21.8 21.8 22.4 24.9 26.2

Zr 140 164 131 127 117 110 115 110 173 175

Nb 9.47 12.7 11.2 15.5 10.7 11.0 13.4 11.8 16.5 18.5

Ba 296 257 113 244 289 233 280 250 134 134

La 17.9 21.1 15.3 17.9 16.0 14.3 16.5 14.4 15.4 15.4

Ce 37.5 46.5 32.6 38.4 34.0 31.5 35.3 30.7 38.8 38.1

Pr 4.87 5.86 4.16 4.99 4.40 4.09 4.67 4.13 5.54 5.42

Nd 20.0 24.4 17.4 21.2 18.3 17.4 19.7 17.8 25.1 24.8

Sm 5.05 6.06 4.49 5.00 4.40 4.26 4.70 4.35 6.22 6.17

Eu 1.46 1.72 1.38 1.89 1.46 1.54 1.78 1.64 2.07 2.05

Gd 5.52 6.58 5.02 5.06 4.68 4.47 4.84 4.61 6.02 6.29

Tb 0.92 1.11 0.87 0.80 0.78 0.73 0.76 0.74 0.93 0.95

Dy 5.77 6.91 5.48 4.69 4.75 4.40 4.49 4.48 5.25 5.27

Ho 1.24 1.47 1.18 0.96 1.01 0.92 0.92 0.92 1.00 0.98

Er 3.33 4.00 3.21 2.49 2.69 2.45 2.39 2.42 2.44 2.50

Tm 0.50 0.59 0.48 0.36 0.40 0.36 0.34 0.34 0.33 0.33

Yb 3.16 3.71 2.96 2.22 2.52 2.23 2.13 2.13 1.96 1.99

Lu 0.48 0.56 0.45 0.34 0.38 0.34 0.32 0.32 0.28 0.28

Hf 3.69 4.27 3.33 2.98 2.89 2.69 2.73 2.67 4.31 4.44

Ta 0.68 0.86 0.78 1.01 0.72 0.72 0.88 0.78 1.31 1.17

Pb 8.22 5.90 5.38 2.33 4.40 1.52 1.98 2.40 1.92 0.21

Th 5.55 5.59 3.81 2.13 3.34 2.48 2.11 1.98 1.22 0.01

U 0.87 0.94 0.66 0.36 0.30 0.37 0.27 0.27 0.44 0.42

Ref.: recommended values from http://georem.mpch-mainz.gwdg.de.


http://georem.mpch-mainz.gwdg.de/

10

Supplementary Table 2 (continued)

Sample BHVO-2 Ref. GSR-1 Ref. AGV-2 Ref. W-2 Ref.


SiO2

TiO2

Al2O3

FeOT

MnO

MgO

CaO

Na2O

K2O

P2O5

L.O.I

Total


Sc 31.9 31.8 5.95 6.1 12.0 13 36.6 36

V 309 318 24.1 24 115 119 265 266

Cr 276 287 3.94 3.6 18.2 17 85.9 92

Co 44.9 44.9 2.99 3.4 15.3 15.5 43.7 43

Ni 108 120 1.26 2.3 18.1 18.9 71.5 72

Cu 125.3 129.3 3.21 3.2 48.5 51.5 105 103

Ga 21.0 21.4 21.5 19 20.7 20.4 18.4 17

Rb 9.2 9.3 463 466 67.4 67.8 19.2 20

Sr 398 394 118 106 627 658 197 197

Y 23.0 25.9 79.1 62 19.8 20 22.5 21.8

Zr 173 171 175 167 225 230 97.1 100

Nb 16.4 18.1 43.9 40 12.6 14.1 6.69 7.51

Ba 136 131 350 343 1086 1134 174 170

La 15.3 15.2 57.9 54 37.6 38 11.0 11.4

Ce 38.1 37.5 110 108 68.3 68 24.0 24

Pr 5.47 5.34 12.9 12.7 8.24 8.3 3.23 3.02

Nd 24.8 24.3 47.1 47 30.8 30.5 13.7 14

Sm 6.25 6.02 10.0 9.7 5.65 5.7 3.41 3.4

Eu 2.07 2.04 0.91 0.85 1.59 1.55 1.13 1.15

Gd 6.08 6.21 9.96 9.3 4.66 4.68 3.82 3.73

Tb 0.93 0.94 1.65 1.65 0.62 0.64 0.63 0.63

Dy 5.24 5.28 11.2 10.2 3.46 3.55 3.92 3.83

Ho 0.99 0.99 2.5 2.05 0.69 0.68 0.86 0.84

Er 2.44 2.51 6.4 6.5 1.77 1.79 2.31 2.17

Tm 0.33 0.33 1.23 1.06 0.26 0.26 0.34 0.33

Yb 1.96 1.99 7.37 7.4 1.64 1.65 2.12 2.1

Lu 0.28 0.28 1.13 1.15 0.26 0.25 0.33 0.33

Hf 4.28 4.47 5.77 6.3 4.86 5.08 2.57 2.6

Ta 1.13 1.15 8.15 7.2 0.87 0.87 0.50 0.5

Pb 1.48 1.65 36.4 31 13.51 13.1 8.44 7.6

Th 1.20 1.22 54.7 54 6.06 6.1 2.26 2.1

U 0.44 0.41 18.3 18.8 1.89 1.885 0.54 0.53

Ref.: recommended values from http://georem.mpch-mainz.gwdg.de. (Fig. 5a and b), the Woniusi lavas were plotted in the subalkaline basalt field with a few transitional into the alkaline basalt and andesite/basalt fields. All the Woniusi basaltic samples fall along the tholeiitic lineon the TiO2 versus FeOT/MgO diagram (Fig. 5c), and show low TiO2 contents (1.2–2.2


http://georem.mpch-mainz.gwdg.de/

11

wt.%; normalized to 100% LOI-free), as Panjal low-Ti basalts do (Fig. 5d).

Fig. 5. (a) Zr/TiO2 versus Nb/Y (Winchester and Floyd, 1976). (b) SiO2 versus Zr/TiO2 (Winchester and Floyd, 1976). (c) TiO2 versus FeOT/MgO. (d) TiO2 versus Mg# (modified from Shellnutt et al. 2015). Data sources for literature data of the Woniusi basalts are from Xiao et al. (2003), Huang et al. (2012), Yu (2013) and Liao et al. (2015)

and the Panjal high-Ti and low-Ti basalts are from Shellnutt et al. (2014).


12

Fig. 6. Chondrite-normalized REE diagrams (a, c, e) and primitive mantle-normalized multi-element variation diagrams (b, d, f). Also shown is a typical ocean island basalt (OIB) and continental arc basalt (CAB) from Sun and McDonough (1989) and Kelemen et al. (2003). Normalization values from Sun and McDonough (1989). Data of the Woniusi basalts are from Xiao et al. (2003), Huang et al. (2012), Yu (2013) and Liao et al. (2015); the Selong basalts are from Zhu et al. (2010); the Qiangtang mafic dykes are from Zhai et al. (2013) and Wang et al. (2014); the Abor mafic volcanics are from Singh and Singh (2012); the Panjal high-Ti and low-Ti basalts are from Shellnutt et al. (2014) and Emeishan low-Ti basalts are from Xu et al. (2001), Xiao et al. (2004) and Wang et al. (2007).

The Woniusi basaltic rocks show a relatively large range of total rare earth elements (REEs;

71.8–131 ppm) and are slightly enriched in LREEs relative to flat HREE [(La/Yb)N = 2.86–5.77, (Dy/Yb)N = 1.21–1.49; Fig. 6a] with weak negative to positive Eu anomalies (δEu = 0.83–1.15). On the primitive-mantle normalized trace element spidergram (Fig. 6b), they are marked by strong enrichment in large ion lithophile elements (LILE) relative to high field strength elements (HFSE). Furthermore, the Woniusi basaltic rocks display well-developed negative Nb–Ta and Sr anomalies, which are similar to those of continental arc basalts (CAB, Fig. 6a and b). 5.3 Whole-rock Sr–Nd–Pb isotopes

The Sr–Nd–Pb isotope analyses are calculated using the ~280 Ma 40

Ar/39

Ar age reported in this study (Supplementary Table 3). The initial

87Sr/

86Sr values are quite high for the Woniusi basaltic rocks

ranging from 0.70451 to 0.70942. The initial 143

Nd/144

Nd ratios range from a low of 0.51203 in sample DSP-17 to a high of 0.51233 in sample WNS-11. The εNd(t) values of the basaltic rocks range between -4.76 and +0.92 (Fig. 7a). The meaningful TDM values for samples with fSm/Nd between -0.28 and -0.16 ranges from 1.21 to 2.08 Ga. The Woniusi basaltic rocks have a rather restricted range in (

206Pb/

204Pb)i

(18.34–18.66), (207

Pb/204

Pb)i (15.68–15.76) and (208

Pb/204

Pb)i (38.72–39). In the Pb-Pb isotope space (Fig. 7b and c), (

207Pb/

204Pb)i and (

208Pb/

204Pb)i positively correlate with (

206Pb/

204Pb)i and plot well

above the Northern Hemisphere Reference Line (NHRL; Hart, 1984). In the 143

Nd/144

Nd versus 206

Pb/ 204

Pb plot (Fig. 7d), the Woniusi basaltic samples slightly deviate from EMII. In general, the most evolved rocks (low MgO) have the lowest εNd(t) and (

206Pb/

204Pb)i values and the highest initial

87Sr/

86Sr ratios of all the rocks.

Fig. 7. Plots of initial Sr, Nd, and Pb isotopic compositions of the Woniusi basalts. In Fig. a, the solid curves represent AFC modeling between the most Nd-depleted Panjal low-Ti basalt sample (PJ1-034) and crustal melts.

The upper crust melt is represented by the Cenozoic Shuangmaidi peraluminous granite (YSZK0-1b) with 63.8 ppm Sr, (87

Sr/86

Sr)i of

0.759, 16.7 ppm Nd, εNd (t=280 Ma) of -10.06 (Huang et al. 2011). The Sr–Nd data of lower crust are from Huang et al. (2007).

Approximate locations of mantle end-members (Zindler and Hart 1986) are indicated for reference. The Northern Hemisphere Reference

Line (NHRL, Hart 1984) is shown. Data sources: the Woniusi basalts from Xiao et al. (2003), Huang Yong et al. (2012), Yu Junchuan

(2013) and Liao et al. (2015); the Daxueshan ultramafic-mafic intrusion from Wang et al. (2018); the Selong basalts from Zhu et al.

(2010); the Qiangtang mafic dykes and flood basalts from Zhai et al. (2013), Wang et al. (2014) and Zhang and Zhang (2017); the Panjal

high-Ti and low-Ti basalts from Chauvet et al. (2008) and Shellnutt et al. (2014); the Emeishan low-Ti basalts are from Xu et al. (2001),

Xiao et al. (2004), Wang et al. (2007) and Xu et al. (2007).


13

Supplementary Table 3 Sr−Nd−Pb isotopes for the Woniusi basalts

Samples 87

Rb

/86

Sr

87Sr

/86

Sr ±2σ

(87

Sr

/86

Sr)i

147Sm

/144

Nd

143Nd

/144

Nd ±2σ

(143

Nd

/144

Nd)i TDM(Ga) ɛNd(t)

206Pb

/204

Pb ±2σ

207Pb

/204

Pb ±2σ

208Pb

/204

Pb ±2σ

(206Pb

/204

Pb)i

(207Pb

/204

Pb)i

(208Pb

/204

Pb)i

Dsp-2 0.0602 0.707395 ± 6 0.707155 0.1649 0.512536 ± 4 0.512234 1.92 -0.85 18.8293 ± 87 15.7805 ± 73 39.5714 ± 184 18.538 15.765 38.855

Dsp-5 0.0573 0.707373 ± 7 0.707145 0.1612 0.512521 ± 4 0.512226 1.83 -1.01 18.8132 ± 13 15.7239 ± 11 39.5718 ± 28 18.545 15.710 38.721

Dsp-10 0.1633 0.705157 ± 6 0.704506 0.1429 0.512584 ± 3 0.512322 1.22 0.87 19.0838 ± 20 15.7012 ± 16 39.8484 ± 42 18.657 15.679 38.925

Dsp-11 0.1367 0.705266 ± 5 0.704721 0.1461 0.512574 ± 3 0.512306 1.30 0.56 19.0463 ± 57 15.7417 ± 46 39.6808 ± 117 18.542 15.716 38.892

Dsp-16 0.1779 0.706967 ± 6 0.706258 0.1539 0.512486 ± 4 0.512203 1.70 -1.45 19.0481 ± 4 15.7480 ± 4 39.6811 ± 17 18.544 15.722 38.785

Dsp-17 0.3487 0.710805 ± 6 0.709416 0.1512 0.512311 ± 4 0.512034 2.05 -4.76 18.8055 ± 15 15.7450 ± 12 39.5175 ± 33 18.432 15.726 38.790

Dsp-18 0.7720 0.712069 ± 6 0.708993 0.1525 0.512314 ± 4 0.512034 2.08 -4.75 18.7464 ± 4 15.7313 ± 3 39.3840 ± 13 18.442 15.715 38.750

Dsp-19 0.1750 0.708116 ± 8 0.707419 0.1500 0.512370 ± 4 0.512095 1.87 -3.56 18.8571 ± 23 15.7468 ± 19 39.6963 ± 50 18.397 15.723 38.802

WNS-11 0.2331 0.706810 ± 6 0.705881 0.1426 0.512586 ± 5 0.512325 1.21 0.92 19.0434 ± 107 15.7416 ± 88 39.8068 ± 226 18.601 15.719 38.943

WNS-16 0.0346 0.705400 ± 5 0.705262 0.1481 0.512524 ± 3 0.512252 1.46 -0.49 19.2472 ± 19 15.7506 ± 15 40.3922 ± 44 18.533 15.714 38.828

WNS-20 0.2057 0.705691 ± 6 0.704871 0.1480 0.512589 ± 4 0.512318 1.30 0.78 18.9519 ± 63 15.7270 ± 52 39.7775 ± 134 18.629 15.710 38.998

The initial isotopic ratios are calculated at 280 Ma, using relevant element concentration measured by ICP-MS.

5.4 Mineral compositions

Chemical compositions of the analyzed clinopyroxenes in three basaltic samples (DSP-11, DSP-12 and WNS-12) are presented on Supplementary Table 4. Clinopyroxenes from Woniusi basalts span a compositional field from diopside to augite (En38.9-48.8Fs12-19.8Wo36.7-45.9 ), with Mg# values from 66.8 to 80.4 [molar MgO/(MgO + FeOT)] and variable contents of TiO2 (0.52–1.61 wt.%), Al2O3 (1.48–4.06 wt.%) and Na2O (0.2–0.37 wt.%). The clinopyroxene phenocrysts all fall within the field of subalkaline (i.e. tholeiitic and calc-alkaline) basalts (Leterrier et al., 1982). 6 Discussion 6.1 Alteration effect

The variably altered nature of the Woniusi basaltic rocks implies that before any petrogenetic and geodynamic inferences can be drawn from the chemistry of the rocks, the possible chemical effects of post-magmatic mobility of elements must be accounted for. The fluid-mobile elements, such as Rb, Ba, Sr and K, show considerably variable (Fig. 8a-d), largely due to an alteration effect. For the Woniusi basaltic rocks, REEs (e.g., La, Ce and Lu), high field strength elements (HFSE, e.g., Nb), U and Th correlate well with Zr (Fig. 8e-j), implying that these elements were essentially immobile during alteration (Pang et al., 2016). Therefore, we conclude that the εNd(t), (

206Pb/

204Pb)i values

and the concentrations of REEs, HFSE, U and Th reported in this study are primary and can be used to evaluate the petrogenetic processes of these basaltic rocks.

6.2 Degrees of crustal contamination


14

Supplementary Table 4 Analyses of clinopyroxene from the Woniusi basalts

Sample WNS-12 DSP-11 DSP-12

Spot 1 2 3 4 5 6 1 2 3 4 1 2 3 4 5 6 7

Type P P P P G G G G G G P P G G G G G


SiO2 51.1 51.6 51.4 51.8 49.2 50.8 50.0 49.1 49.3 48.8 51.7 52.5 50.0 49.1 49.0 49.6 50.3

TiO2 0.52 0.71 0.85 0.99 1.11 1.22 1.35 1.35 1.61 1.29 0.54 0.61 1.50 1.46 1.52 1.25 0.99

Al2O3 2.73 1.44 2.48 2.64 3.61 2.99 3.34 4.06 2.76 3.87 1.82 1.87 3.00 2.84 3.01 3.54 3.40

FeOT 7.57 11.6 9.04 9.27 8.98 10.7 8.89 8.95 11.7 7.89 7.47 7.42 9.38 10.6 9.53 8.23 7.34

MnO 0.22 0.251 0.1 0.113 0.22 0.23 0.18 0.24 0.28 0.20 0.26 0.20 0.26 0.31 0.25 0.20 0.14

MgO 16.6 16.1 16.2 16.0 14.5 14.2 14.2 14.5 13.2 14.6 17.0 17.1 14.7 15.1 15.4 14.9 14.9

CaO 18.2 18.4 20.3 19.7 21.0 19.7 21.8 21.7 19.6 21.6 19.9 20.0 21.3 21.1 21.8 21.5 20.8

K2O 0.03 0.02 0.01 0.00 0.01 0.01 0.01 0.00 0.01

Na2O 0.37 0.16 0.14 0.12 0.21 0.20 0.33 0.32 0.34 0.30 0.21 0.20 0.27 0.25 0.28 0.27 0.26

NiO 0.43 0.01 0.38 0.08 0.02 0.18 0.04 0.05 0.13 0.19 0.21 0.02

Cr2O3 1.07 0.06 0.43 0.30 0.29 0.07 0.33 0.43 0.07 0.65 0.59 0.72 0.24 0.02 0.03 0.38 0.60

Total 98.9 100.3 101.0 101.0 99.1 100.5 100.4 100.7 99.0 99.2 99.7 100.7 100.7 100.9 101.0 100.0 98.7

Cations (O=6)

Si 1.90 1.92 1.88 1.90 1.84 1.89 1.85 1.81 1.88 1.82 1.91 1.92 1.85 1.81 1.80 1.84 1.88

Ti 0.01 0.02 0.02 0.03 0.03 0.03 0.04 0.04 0.05 0.04 0.01 0.02 0.04 0.04 0.04 0.03 0.03

Al 0.12 0.06 0.11 0.11 0.16 0.13 0.15 0.18 0.12 0.17 0.08 0.08 0.13 0.12 0.13 0.15 0.15

Fe3+

0.04 0.07 0.08 0.03 0.10 0.03 0.09 0.13 0.06 0.11 0.07 0.04 0.10 0.19 0.20 0.11 0.03

Fe2+

0.19 0.29 0.20 0.26 0.19 0.30 0.19 0.14 0.31 0.13 0.16 0.19 0.19 0.14 0.09 0.15 0.20

Mn 0.01 0.01 0.00 0.00 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.00

Mg 0.92 0.89 0.89 0.87 0.81 0.79 0.78 0.80 0.75 0.81 0.93 0.93 0.81 0.83 0.84 0.82 0.83

Ca 0.73 0.73 0.80 0.78 0.84 0.79 0.87 0.86 0.80 0.86 0.79 0.78 0.84 0.83 0.86 0.85 0.83

K 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Na 0.03 0.01 0.01 0.01 0.02 0.01 0.02 0.02 0.03 0.02 0.02 0.01 0.02 0.02 0.02 0.02 0.02

Ni 0.01 0.00 0.00 0.00 0.00 0.01 0.00 0.00 0.00 0.00 0.01 0.00 0.00 0.00 0.01 0.01 0.00

Cr 0.03 0.00 0.01 0.01 0.01 0.00 0.01 0.01 0.00 0.02 0.02 0.02 0.01 0.00 0.00 0.01 0.02

Total 4.00 4.00 4.00 4.00 4.00 4.00 4.00 4.00 4.00 4.00 4.00 4.00 4.00 4.00 4.00 4.00 4.00

End-member compositions (mol%)

En 48.8 44.8 45.1 45.1 41.7 41.2 40.6 41.0 38.9 42.1 47.6 47.8 41.4 41.5 42.0 42.4 43.8

Fs 12.8 18.5 14.3 14.9 14.9 17.8 14.6 14.6 19.8 13.1 12.2 12.0 15.3 16.8 15.0 13.5 12.3

Wo 38.4 36.7 40.6 40.1 43.4 41.0 44.9 44.3 41.4 44.8 40.2 40.3 43.3 41.7 43.0 44.1 43.9

P, phenocryst; G, groundmass

Variable degrees of crustal assimilation are commonly expected for mantle-derived magmas

migrating through the continental crust to the surface (DePaolo, 1981; Spera and Bohrson, 2001). The correlations between indices of fractionation and trace element/isotopic ratios in basaltic rocks seem to provide good evidence for crustal contamination (Pang et al., 2016). Together with published data, most of the Woniusi basaltic rocks have small variations of εNd(t) values, showing no significant correlations with MgO, Nb/La and Pb/Ce (Fig. 9a-c), ruling out significant crustal contamination during magma ascending. On the (

206Pb/

204Pb)i versus SiO2 diagram (Fig. 9d), no obvious correlation

can be recognized and (206

Pb/204

Pb)i ratios are relatively constant with the decreasing SiO2, suggesting that fractional crystallization might have exerted a dominant control on the magma process of the these lavas rather than fractionation crystallization combined with crustal assimilation (AFC). Therefore, these samples were not affected by crustal contamination and their trace elements and isotopic compositions could approximately reflect the characteristics of their mantle sources.

In contrast, the remaining samples of the Woniusi basaltic rocks show positive trends in the plots of εNd(t) versus MgO and Nb/La (Fig. 9a and b), and negative trends in the plots of εNd(t) versus Pb/Ce and (

206Pb/

204Pb)i versus SiO2 (Fig. 9c and d). It is apparent that these lavas may have undergone

appreciable crustal contamination. The amount of crustal contamination can be constrained by an AFC process involving the most depleted Panjal low-Ti sample (PJ1-034) and the Baoshan upper/lower crust. As shown in Fig. 7a, the proportion of crustal input is estimated to be ~10−15% for the samples with highly isotopically enriched characteristics (εNd(t) = -4.76 to -1.45).


15

Fig. 8. Bivariate trace element plots to determine the extent of correlation of various immobile and mobile trace elements. Literature data of the Woniusi basalts are from Xiao et al. (2003), Huang et al. (2012), Yu (2013) and Liao et al. (2015). 6.3 Nature of mantle source

A key feature of the Woniusi basalts is the LILE (e.g., Th and U) and LREE enrichment and the presence of negative Nb and Ta anomalies (Fig. 6a and b), which contrasts with the signature of


16

plume-related ocean island basalts (OIB, Fig. 6a and b). These geochemical characteristics are well-established features for major continental flood basalts (Arndt and Christensen, 1992; Arndt et al., 1993; Puffer, 2001; Ewart et al., 2004; Jourdan et al., 2007; Wang et al., 2007; Wei et al., 2014) and are generally interpreted as reflecting either (1) a mantle source composition modified by metasomatic enrichment induced by silicate melt or hydrous fluid percolation, possibly reflecting ancient subduction (Puffer, 2001; Ewart et al., 2004; Jourdan et al., 2007; Wang et al., 2007; Wei et al., 2014), (2) asthenosphere- or plume-derived magmas interacting with lithospheric mantle (Arndt and Christensen, 1992; Jourdan et al., 2007), or (3) crustal assimilation by mantle-derived magmas (Arndt et al., 1993). In all three cases, the negative Nb and Ta anomalies of the continental flood basalt magmas would have been inherited within the lithosphere (Jourdan et al., 2007). For most of the Woniusi basaltic samples, as discussed above, crustal contamination appears to be in most cases negligible and thus cannot account for the systematically observed negative Nb and Ta anomalies. If we accept this feature to be diagnostic of the mantle source, then an sub-continental lithospheric mantle (SCLM) source metasomatized by silicate melts or fluids is implied. Moreover, the Sr–Nd–Pb isotopic compositions of the Woniusi basalts are characterized by relatively low εNd(t) values and high

87Sr/

86Sr and

206Pb/

204Pb

isotope ratios, placing them close to the EMII pole (Fig. 7; Zindler and Hart, 1986; Workman et al., 2004; Workman and Hart, 2005). The depleted mantle model ages (TDM) of the Woniusi samples with fSm/Nd between -0.28 and -0.22 range from 1.21 to 2.08 Ga and are concordant with Nd model ages (1.73–2.16 Ga) of the Early Paleozoic and Mesozoic to Cenozoic granites from the Baoshan terrane (Chen et al., 2007; Huang et al., 2011), underscoring a need for long-term preservation of such ancient (Proterozoic) enriched signatures in the SCLM source. However, considering that the Sr–Nd isotope signatures of most of the uncontaminated Woniusi basalts are depleted relative to present-day Bulk Earth (Supplementary Table 3), it seems reasonable to conclude that the mantle enrichment age does not appear to be very old, such as Paleoproterozoic, and the metasomatic enrichment of their source region which may occur relatively recently, possibly at Early Paleozoic (Li et al., 2015).

Fig. 9. Plots of εNd(t) versus MgO (a), εNd(t) versus Nb/La (b), εNd(t) versus Pb/Ce (c) and (206Pb/204Pb)i versus SiO2 (d) to assess the effect of crustal assimilation on the compositions of the Woniusi basalts. Dashed outlined fields highlight that these samples display trends parallel to that of the fractional crystallization and thus are not affected by shallow-level crustal assimilation.

The continental upper mantle is dominated by peridotites containing olivine, orthopyroxene,

clinopyroxene and an aluminous phase such as spinel or garnet (Pang et al., 2015). The modelling of the upper mantle partial melting processes can be illustrated using the plots of Dy/Yb versus La/Yb


17

ratios; such plots can also distinguish between melting in the spinel and garnet stability fields (e.g., Jung et al., 2012; Huang et al., 2013; Ayalew et al., 2016). An additional benefit of such plots is that mixing of melts from distinct mantle sources produces linear mixing arrays. To constrain the degree of partial melting, we perform REE modeling on spinel and garnet peridotite using the non-modal batch melting equations of Shaw (1970). The results indicate that ~4–10% melting of the spinel lherzolite can generate the observed variation in La/Yb, but fail to explain the variation in Dy/Yb (Fig. 10a). Thus, to account for the small range in Dy/Yb ratios (1.81–2.23) observed in the Woniusi samples, melts from garnet-facies mantle (5%–6%) and spinel-facies mantle (4%–5%) have to mix. When plotted on the Yb versus La/Yb diagram (Fig. 10b), a similar result can also be obtained. This indicates that the partial melting could have occurred at the spinel-garnet transition zone, corresponding to a depth of 60−85 km (McKenzie and O'Nions, 1991; Robinson and Wood, 1998). Therefore, partial melting events to producing the Woniusi basalts mainly constricted within depth of 60–85 km. Such a constraint from fractionation of REE agrees well with the present-day lithosphere thickness of the southwestern Yunnan obtained by S-wave receiver functions (80−100 km; Hu et al., 2012).

Substantial melting, as would be required to produce continental flood basalts, of the SCLM poses a problem if conditions are dry (Arndt and Christensen, 1992), but is likely to occur if previous hydration by volatiles had taken place (Gallagher and Hawkesworth, 1992; Turner et al., 1996). The ancient SCLM under Baoshan is inevitably modified by subduction-related processes (melt or fluid metasomatism) (Ben Othman et al., 1989; Noll Jr. et al., 1996; Class et al., 2000; Johnson and Plank, 2000; Wei et al., 2014) during the Early Paleozoic continental accretion (Wang et al., 2013).The uncontaminated Woniusi basaltic samples plot within the feld of global continental arc basalts in the Th/Yb versus Nb/Yb diagram (Fig. 10c), consistent with magma generation related to subduction metasomatism. Contribution of sediments to the mantle source of the Woniusi lavas is evident on the Sr/La versus La/Yb (Fig. 10d) and Th/Nb versus Pb/Ce (not shown) diagrams, as these variables are reliable indicators of potential sediment or fluid contributions to magma source regions (Woodhead et al., 2001). Therefore, the lithospheric mantle domains under Baoshan were likely enriched by recycled sediments, which may have imparted an EMII-like signature (Callegaro et al., 2014).

Fig. 10. Plots of Dy/Yb (a), Yb (b) versus La/Yb, Th/Yb versus Nb/Yb (c) and Sr/La versus La/Yb (d) for the uncontaminated Woniusi basalts. Non-modal batch melting curves were calculated using partition coefficients from McKenzie and O’Nions (1991). Blue curves: melting

curve of a spinel-bearing lherzolite source (modal composition and melt mode of 53% olivine, 27% orthopyroxene, 17% clinopyroxene,

3% spinel and 6% olivine, 28% orthopyroxene, 67% clinopyroxene, 11% spinel, respectively; Kinzler 1997). Pink curves: melting curve

of a garnet-bearing lherzolite source (60% olivine, 20% orthopyroxene, 10% clinopyroxene, 10% garnet and 3% olivine, 3%


18

orthopyroxene, 70% clinopyroxene, 24% garnet, respectively; Walter 1998). The dashed lines represent mixing curves between melt

derived from 5% melting of a source with 3% residual spinel and from 5% melting of a source with 10% residual garnet. The assumed

source of non-modal batch melting curves is ‘‘99% primitive mantle (McDonough and Sun 1995) + 1% upper crust (Rudnick and Gao

2003)”. The values of OIB, CAB, and MORB for comparison are from Wang et al. (2018). Literature data sources are the same as in Fig.

5.

6.4 The Early Permian magmatism on the northern margin of East Gondwana Previous paleontologic analyses of conodonts in limestone intercalations of the Woniusi Formation

reported a Late Artinskian to Early Kungurian age (287–272 Ma), which is interpreted to represent the eruption age for the Woniusi magmatism (Wang et al., 2004). However, Wang et al. (1999) doubted that the limestone intercalations of the Woniusi Formation were faulted repetitions of the underlying Dingjiazhai Formation; therefore, the correlated age is likely to be questionable. Moreover, the Woniusi Formation is paraconformably underlain by the Dingjiazhai Formation, which is defined as siltstone, glacio-marine diamictite, pebbly mudstone, dark shale and bioclastic limestone (Jin et al., 2011). Fusulinoideans in limestone units within the top part of the Dingjiazhai Formation, such as Sweetognathus bucaramangus (Rabe), S. whitei (Rhodes), and Mesogondolella bisselli (Clark and Behnken) reveals that the timing of Woniusi volcanic activity is later than middle Artinskian (~287 Ma; Ueno et al., 2002).

It is notable that the Woniusi basalts in the Xiyi area contain zircons, some of which exhibit features typical of a magmatic origin. LA–ICP–MS U–Pb dating of these magmatic zircons reported an age of 261.3 ± 1.2 Ma interpreted to represent the formation age of their host basalts (Yao, 2016). Although several studies have suggested that zircon could crystallize from basaltic magmas (e.g., Tietz and Büchner, 2007; Hurai et al., 2010), most of the crystals observed in the basalts are identified as xenocrysts, derived from either crustal or mantle rocks (e.g., Belousova et al., 2002; Siebel et al., 2009; Li et al., 2014). In this case, the zircon age is likely to be questionable. Because previously estimated ages of Woniusi Formation were poorly defined, it is difficult to evaluate their geodynamic importance of this episode of magmatism. Our new

40Ar/

39Ar age data suggest that the Woniusi basaltic volcanism

in the Baoshan terrane were formed in Early Permian times (280–274 Ma; Fig. 4). Notably, the 40

Ar/ 39

Ar age is compatible with the primary magmatic zircon LA–ICP–MS U–Pb ages of 282.4 ± 2.5 Ma and 300.9 ± 2.0 Ma for two diabase dykes in the same region, which constrain the emplacement ages of the cogenetic mafic dykes (Liao et al., 2015). Recently, SHRIMP U-Pb dating of zircon crystals from the gabbro for the Daxueshan mafic–ultramafic intrusion in northern Baoshan, which is exposed surrounding the Woniusi basalts and diabase dykes, yields a crystallization age of 300.5 ± 1.6 Ma (Wang et al., 2018).

Early Permian magmatism is also present in the northern margin of the Indian continent. The rhyolite from the lower-middle volcanic sequence of the Panjal Traps in the India Himalaya has yielded a zircon U–Pb age of 289 ± 3 Ma (Shellnutt et al., 2011; Supplementary Table 5) and the Yunam granitic dykes in the Indian High Himalaya provided a zircon U–Pb age of 284 ± 1 Ma (Spring et al. 1993). To the south, the Bhote Kosi-Selong basalts formed at ~280 Ma constraint by brachiopods in the associated sedimentary sequences have been reported in the Tethyan Himalaya (Garzanti et al., 1999; Zhu et al., 2010). Most recently, a paleomagnetic study concluded that the Abor volcanics in the lesser Himalaya were emplaced in the Early Permian (Ali et al., 2012). Early Permian magmatism is extensively developed in the adjacent continental terranes as well, including South Qiangtang (285–279 Ma) and Lhasa (~280 Ma) (Supplementary Table 5, Zhu et al., 2010; Zhai et al., 2013).

Fig. 11. Compilation of reliable 40Ar/39Ar and zircon U−Pb ages of basalts, mafic dykes and ultramafic-mafic intrusion from the Woniusi, Panjal Traps and South Qiangtang. The numbers at the peak of probability curves (ca. 283 Ma) represent the averaged ages of both literature data and those reported in this study. The data sources are given in Supplementary Table 5.


19

Supplementary Table 5 A summary of the Early Permian magmatism along the northern margin of East Gondwana

Terrane Flood basalt province Coverage (km2) Stage Locality and lithology Age (Ma) Method References

Baoshan

Woniusi 12000 Middle-late Artinskian–Kungurian Paleontology Wang Wei et al., 2004

Daxueshan, ultramafic-mafic

intrusion 300.5 ±1.6 SHRIMP U–Pb zircon Wang et al., 2018

Jinji, mafic dyke 282.4 ± 2.5 LA–ICPMS U–Pb zircon Liao et al., 2015

Jinji, mafic dyke 300.9 ± 2.0 LA–ICPMS U–Pb zircon Liao et al., 2015

Dongshanpo, basalt 279.5 ± 1.4 Groundmass 40

Ar/39

Ar This study

Jinji, basalt 273.9 ± 1.5 Groundmass 40

Ar/39

Ar This study

Qiangtang

Qiangtang 40000 Sakmarian–Kungurian Paleontology Zhang and Zhang, 2017

Northern Gangmacuo, mafic dyke 284 ± 3 SHRIMP U–Pb zircon Zhai Qingguo et al., 2009

Lugu, mafic dyke 302 ± 4 SHRIMP U–Pb zircon Zhai Qingguo et al., 2009

Southern Gangmacuo, mafic dyke 282.8 ± 1.4 SHRIMP U–Pb zircon Zhai et al., 2013

Gemuri, mafic dyke 279.0 ± 1.6 SHRIMP U–Pb zircon Zhai et al., 2013

Gangmacuo, mafic dyke 285.1 ± 1.4 SHRIMP U–Pb zircon Zhai et al., 2013

Gangmacuo, mafic dyke 285.0 ± 2.5 SHRIMP U–Pb zircon Zhai et al., 2013

Gangmacuo, mafic dyke 291 ± 2 LA–ICPMS U–Pb zircon Wang et al., 2014

Gangmacuo, mafic dyke 292 ± 3 LA–ICPMS U–Pb zircon Wang et al., 2014

Taoxinghu, mafic dyke 300 ± 1.8 LA–ICPMS U–Pb zircon Wang et al., 2014

Hongjishan, mafic dyke 290.6 ± 3.5 LA–ICPMS U–Pb zircon Xu et al., 2016

Hongjishan, mafic dyke 290.1 ± 1.5 LA–ICPMS U–Pb zircon Xu et al., 2016

Zaduo, peridotite 275.3 ± 1.9 Hornblende 40Ar/39Ar Wang Yizhi et al., 2007

Ningduo, basalt 272.9 ± 0.9 LA–ICPMS U–Pb zircon Zhou Huiwu, 2013

Zaduo, mafic dyke 284 ± 2 LA–ICPMS U–Pb zircon Liu Bin, 2014


20

Northern India

Panjal Traps 10000 Sakmarian−Artinskian Paleontology Shellnutt et al., 2015

Panjal Traps, rhyolite 289.0 ± 2.5 LA–ICPMS U–Pb zircon Shellnutt et al., 2011

Yunam, granitic dyke 284 ± 1 TIMS U–Pb zircon Spring et al., 1993

Abor 2500 Artinskian Paleomagnetism Ali et al., 2012

Bhote Kosi-Selong Unkown Late Sakmarian−Kungurian Paleontology Garzanti et al., 1999

Lhasa Jiangrang Unkown Sakmarian–Artinskian

Paleontology Zhu et al., 2010


21

Fig. 11 shows that that the widespread Early Permian magmatism was formed contemporaneously at

ca. 283 Ma. Additionally, recent paleomagnetic and paleontological reconstructions suggest that the Baoshan terrane was geographically in the vicinity of the northeast Greater India and northwest Australia up until the Early Permian, although there is debate regarding the exact position of the Baoshan terrane (Ali et al. 2013; Metcalfe 2013). The age data, in combination with the spatial distribution of the Early Permian magmatism during the initial rifting (cf. Liao et al., 2015), indicate that the Early Permian magmatism on the northern margin of East Gondwana is temporally and spatially related to each other and could be the products of the same tectono-magmatic event.

6.5 Lithospheric and sub-lithospheric mantle melting above a common mantle plume

Although the conceptual framework of Early Permian rifting related to the opening of a new ocean basin (Meso-Tethys) on the northern margin of East Gondwana seems straightforward (e.g. Sengor, 1987; Stampfli and Borel, 2002; Ali et al., 2012, 2013; Metcalfe, 2013), the cause of the formation and propagation of the rift has remained enigmatic. Two major models have been proposed in literature to trigger the rifting on the northern margin of Gondwana, involving lithospheric extension in the passive continental rifting and mantle plume (e.g. Zhu et al., 2010; Zhai et al., 2013; Shellnutt et al., 2014, 2015; Liao et al., 2015; Wang et al., 2014; Xu et al., 2016; Zhang and Zhang, 2017). The lithospheric extension model requires a stress field to trigger the initiation of the continental rift as a result of lithospheric thinning and passive upwelling of the asthenosphere induced by slab-pull on the far side of the Paleo-Tethys or slab-rollback on the south side of Australia (Shellnutt et al., 2014, 2015). This geodynamic scenario seems likely for the generation of the widespread Early Permian magmatism within the Himalaya. However, the modern lithosphere in the southwestern Yunnan has a thicknesses of ~80−100 km (Hu et al., 2012). Thus, it seems unlikely that large-scale thinning of the lithospheric mantle took place during the Early Permian. Additionally, in accordance with the model of McKenzie and Bickle (1988), the composition and volume of melts is directly related to the amount of lithospheric extension and the potential temperature of the underlying asthenosphere. Therefore, the large volume of Permian magmatism in the Baoshan terrane would require a relatively large amount of lithospheric extension above a mantle of normal potential temperature. However, there is no evidence of widespread Early Permian faulting in the Baoshan terrane. Hence, lithospheric extension apparently was not a likely mechanism for the generation of the large volume of the Woniusi flood basalt province over a short period of time (~20 Ma; Liao et al., 2015).

Another model that could promote continental ruptures and LIP magmatism is a mantle plume activity according to Bryan and Ferrari (2013) and Xu et al. (2015), since there is no strong argument to support lithospheric extension. Recently, the model of mantle plume activity associated with the Early Permian rifting of the Cimmeria from the Indian margin of Gondwana is gaining popularity (Chauvet et al., 2008; Zhu et al., 2010; Zhai et al., 2013; Wang et al., 2014; Liao et al., 2015; Xu et al., 2016; Zhang and Zhang, 2017). In this model, the plume would rise from the base of the lithosphere and expand laterally to form a head that could attain more than 1000 km in diameter (e.g. Richards et al., 1989; Campbell and Griffiths, 1990; Campbell, 2005). Such a mechanism could explain the large-scale (~5.25 × 10

4 km

2; Liao et al., 2015) eruption of the Early Permian basaltic rocks in the

Panjal Traps, Abor, Selong and Qiangtang, and possibly initiate the rifting along the northern Indian margin of East Gondwana (Chauvet et al., 2008; Zhu et al., 2010; Zhai et al., 2013; Wang et al., 2014; Xu et al., 2016; Zhang and Zhang, 2017). Similarly, Liao et al. (2015) have suggested that the Woniusi flood basalt province is correlative to the synchronous basaltic rocks on the northern margin of the Indian continent and that it may represent the northeastern extension of a fragmented LIP with an original extent of ˃ 2 × 10

6 km

2. This Panjal-Qiangtang-Woniusi LIP was most likely the product of an

Early Permian mantle plume on the northern margin of East Gondwana (e.g. Liao et al., 2015; Zhang and Zhang, 2017).

The evidence for a mantle plume model is that the duration of volcanism appears to be very short as the rocks are constrained to the Sakmarian-Kungurian (Supplementary Table 5). The appearance of giant radial dyke swarms is another major characteristic of mantle plume (Bryan and Ernst, 2008). Such dyke swarms occurred in the South Qiangtang and Baoshan terrane (Wang et al., 2014; Liao et al., 2015). Moreover, pre-volcanic lithospheric uplift is ranked as one the most important criteria used to identify the presence of plumes (Ernst and Buchan, 2001; He et al., 2003; Saunders et al., 2007). Liao et al. (2015) indeed speculated a Serpukhovian to Gzhelian (331 Ma to 304 Ma) uplift in the northern Sibumasu terrane based on the Upper Carboniferous erosional unconformities and rift-related sedimentation. The timing of this uplift predated the emplacement of the ~283 Ma Woniusi flood basalt province. All these observations are consistent with the significant role of a mantle plume in the generation of the Early Permian magmatism within the Himalaya.

As noted earlier, a key feature of the Woniusi, Selong and Panjal basalts is their LILE enrichment, the presence of negative Nb and Ta anomalies and negative εNd(t) (-6.02 to +1.35) (Fig. 6 and 7). These signatures argue that these basalts were formed via crystal fractionation with variable crustal material


22

assimilation of basaltic magma derived from the SCLM (Shellnutt et al., 2014; this study). In contrast, the Qiangtang high-Ti mafic dykes and related flood basalts have variable isotope compositions [εNd(t) = +2.39 to +6.69, (

206Pb/

204Pb)i = 18.22 to 19.37] (Xu et al., 2016; Zhang and Zhang, 2017) and display

OIB-like trace element signatures (Fig. 6). These rocks are interpreted as decompression melting within a convecting mantle plume. Geochemical comparisons between the flood basalt provinces from the Baoshan, Tethyan Himalaya, Panjal Traps and from the Qiangtang therefore illustrate the two different mantle domains (i.e., lithospheric and sub-lithospheric mantle) for the coevally emplaced basaltic rocks within the Himalaya.

Fig. 12 illustrates our preferred geodynamic model for the formation of the magmatism in the Early Permian. (a) The regional cold and refractory SCLM underneath East Gondwana was modified and enriched by subduction components over the course of Early Paleozoic. The ca. 501-492 Ma rhyolites and granitoids in the Lhasa terrane,which is not far from the study area, are inferred to be products of the subduction of the proto-Tethyan ocean lithosphere beneath the Australian margin of East Gondwana (Zhu et al., 2013). The subduction would have caused back-arc extension and asthenospheric upwelling. As a result, the subduction-induced decompressional mantle melting would lead to the generation of volcanic rocks within the Gongyanghe Group in the Baoshan terrane (Yang Xuejun et al., 2012). Metasomatized garnet pyroxenite or eclogite veins may have been formed within the SCLM. (b) When the ascending Tarim plume impinged on the enriched SCLM during the Early Permian, conductive heating from the deep-seated plume caused melting of the premetasomatized domains within the SCLM and giving rise to the basalts. Unlike the Woniusi, Selong and Panjal basalts, the Qiangtang mafic dykes and flood basalts were probably derived from a sub-lithospheric mantle source (Wang et al., 2019). Differently, the Qiangtang low-Ti mafic dykes have undergone continental crustal assimilation en route to the surface (Wang et al., 2019). Thus, all these basaltic rocks in the Himalaya could have genetically been related to a Permian upwelling mantle plume. The arrival of the plume head also led to thermal uplift, resulting in lithospheric extension along the northern margin of East Gondwana.

Fig. 12. Schematic cartoon illustrating the interaction betweenan ascending mantle plume with previously metasomatized SCLM and the origin of the ca. 283 Ma Panjal-Qiangtang-Woniusi LIP. Arrows denote the direction of the breakup stresses probably linked to the mantle plume.


23

6.6 Late Paleozoic magmatism related to mantle plume activity in the margin of Pangea supercontinent

During the Late Paleozoic, the Gondwana continent amalgamated with the Laurasia continent, and the Pangea supercontinent was formed (e.g. Metcalfe, 2013; Torsvik and Cocks, 2013; Safonova and Maruyama, 2014). There are many continental LIP events occurred in the margin of Pangea supercontinent throughout the Late Palaeozoic (~290 Ma to ~251 Ma), including the Tarim LIP (~290–280 Ma, Zhang et al., 2013; Xu et al., 2014; Cheng et al., 2018), Emeishan LIP (~260 Ma; Zhang et al., 2006, 2008; Shellnutt et al., 2012; Zhong et al., 2014) and Siberian Traps (~251 Ma, Kamo et al., 2003) (Fig. 13). Obviously, the Panjal-Qiangtang-Woniusi LIP likely occurred ca. 23 Ma before the Emeishan LIP and 32 Ma before the Siberian Traps, but was contemporaneous with the eruption of the Tarim LIP. Based on numerical experiments, geophysical (e.g., high resolution seismic tomography) and geochemical observations, many, if not all, thermal mantle plumes closely associated, in space and time, with some deep slab subduction systems, imply a close genetic connection between slab subduction and formation of some hot mantle plumes (Wang et al., 2013; Cheng et al., 2018). Such hypothesis, which emphasized the importance of slab avalanches in the formation of some hot mantle plumes and the resultant large igneous provinces (LIPs), could also imply that mantle plumes are intimately linked to plate tectonics (Steinberger and Torsvik, 2012). Hence, such a sudden fare up of plume activities (including Tarim, Emeishan and Siberian) in the Permian is attributed to the early stage of the dipolar Pangean and SW Pacific superplumes resulted from circum-Pangea subduction and slab avalanches (Li and Zhong, 2009; Zhang et al., 2010). In view of the location of Sibumasu terrane and Indian continent in the margin of Pangea supercontinent during the Late Paleozoic time (Fig. 13, cf. Ali et al., 2013; Liao et al., 2015; Shangguan et al., 2016), we believe that the ca. 283 Ma Panjal-Qiangtang-Woniusi LIP may be also linked to the above-mentioned subduction.

Fig. 13 Palaeogeographical reconstruction at the Late Paleozoic showing simplified plate boundaries and labels of some major features (modified after Ali et al., 2013; Liao et al., 2015; Shangguan et al., 2016). The dark areas show the recognized LIPs between 290 and 251 Ma. 7 Conclusions

From the results presented in this study the following conclusions can be drawn: (1) The basalts of the Woniusi Formation formed at 280 to 274 Ma, coeval with the basaltic rocks

from the Panjal Traps, Tethyan Himalaya (Abor, Bhote Kosi−Selong), Lhasa, and Qiangtang. (2) The Woniusi basalts are depleted in HFSE (Nb, Ta, Zr, Hf, and Ti) and have low εNd(t) and high

initial Pb isotopes, and were likely generated by partial melting of sub-continental lithospheric mantle that had been metasomatized by previous subduction processes.

(3) The Woniusi flood basalts are likely the remnants of a ca. 283 Ma plume-induced LIP that played a key role in the Early Permian rifting on the northern margin of East Gondwana, and this LIP may be related to circum-Pangea subduction.


24

Acknowledgements

We are grateful to Prof. Zhaochong Zhang and Hongyan Li for their very useful and detailed constructive journal reviews that helped improve our arguments. Thanks are given to Ying Liu, Guangqian Hu, Xianglin Tu and Xiujuan Bai for laboratory assistance. This research was supported by the Geological Survey Project of China Geological Survey (DD20190305), National Natural Science Foundation of China (NSFC) (No. 41703030), China Scholarship Council (No. 201808360273), and the research grant of Shandong Key Laboratory of Depositional Mineralization & Sedimentary Minerals, Shandong University of Science and Technology (No. DMSM20190029).

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About the first author CAO Jun was born in Huanggang of China in 1987, and received his Ph.D. from the University of Chinese

Academy of Sciences in 2015. He is currently a lecturer at the East China University of Technology. His research interests focus on the metallogenesis of Tarim large igneous province. Email: [email protected]; phone: 0791-83897549, 13207080820.


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