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Volume 30 Issue 5
Oct.  2019
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Early-Middle Triassic Intrusions in Western Inner Mongolia, China: Implications for the Final Orogenic Evolution in Southwestern Xing-Meng Orogenic Belt

  • The end-Permian to Early-Middle Triassic magmatic rocks in Inner Mongolia can provide valuable insights into the relationships between the collisional processes and the magmatic responses during the final orogenic evolution of Xing-Meng orogenic belt (XMOB). This paper presents zircon U-Pb ages and Hf isotopes, whole rock geochemical and Sr-Nd-Pb isotopic data for the Early-Middle Triassic diabases and monzogranites from the Langshan area, southwestern XMOB. Our results suggest that the studied diabases and monzogranites were respectively formed during Early Triassic and Middle Triassic. The Early Triassic diabases are characterized by "arc-like" geochemical signatures, including enrichment in Rb, U and K, and depletion in Nb, Ta, P and Ti. They have negative to weak positive εNd(t) values (-3.1 to +1.5) and relatively high initial ratios of 208Pb/204Pb (35.968-37.346), 207Pb/204Pb (15.448-15.508) and 206Pb/204Pb (16.280-17.492), indicating a subduction-metasomatized enriched lithospheric mantle source. Their low Ba/Rb (2.72-6.56), Ce/Y (0.97-1.39) and (Tb/Yb)N ratios (1.31-1.45) suggest that the parental magma was likely originated from low degree partial melting of the phlogopite-bearing lherzolite in a spinel- stability field. The Middle Triassic monzogranites show high Sr/Y ratios, low MgO, Cr and Ni contents, high Zr/Sm ra-tios (40-64), negative zircon εHf(t) values (-25.8 to -8.8), as well as relatively flat heavy rare earth element patterns. They were likely derived from low degree partial melting of a moderately thickened ancient lower crust. The diabases and the slightly postdated high Sr/Y granites in this study represent the magmatic responses to the final orogenic evolution in the southwestern XMOB. To-gether with regional works, we propose that the slab break-off of the Paleo-Asian oceanic lithosphere following the terminal collision between the North China Craton and the South Mongolia terranes triggered asthenospheric upwelling, and the ongoing convergence further initiated moderately crustal thickening and uplift in the XMOB.
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Early-Middle Triassic Intrusions in Western Inner Mongolia, China: Implications for the Final Orogenic Evolution in Southwestern Xing-Meng Orogenic Belt

    Corresponding author: Shaocong Lai, shaocong@nwu.edu.cn
  • 1. State Key Laboratory of Continental Dynamics, Department of Geology, Northwest University, Xi'an 710069, China
  • 2. School of Earth Sciences and Resources, China University of Geosciences, Beijing 100083, China

Abstract: The end-Permian to Early-Middle Triassic magmatic rocks in Inner Mongolia can provide valuable insights into the relationships between the collisional processes and the magmatic responses during the final orogenic evolution of Xing-Meng orogenic belt (XMOB). This paper presents zircon U-Pb ages and Hf isotopes, whole rock geochemical and Sr-Nd-Pb isotopic data for the Early-Middle Triassic diabases and monzogranites from the Langshan area, southwestern XMOB. Our results suggest that the studied diabases and monzogranites were respectively formed during Early Triassic and Middle Triassic. The Early Triassic diabases are characterized by "arc-like" geochemical signatures, including enrichment in Rb, U and K, and depletion in Nb, Ta, P and Ti. They have negative to weak positive εNd(t) values (-3.1 to +1.5) and relatively high initial ratios of 208Pb/204Pb (35.968-37.346), 207Pb/204Pb (15.448-15.508) and 206Pb/204Pb (16.280-17.492), indicating a subduction-metasomatized enriched lithospheric mantle source. Their low Ba/Rb (2.72-6.56), Ce/Y (0.97-1.39) and (Tb/Yb)N ratios (1.31-1.45) suggest that the parental magma was likely originated from low degree partial melting of the phlogopite-bearing lherzolite in a spinel- stability field. The Middle Triassic monzogranites show high Sr/Y ratios, low MgO, Cr and Ni contents, high Zr/Sm ra-tios (40-64), negative zircon εHf(t) values (-25.8 to -8.8), as well as relatively flat heavy rare earth element patterns. They were likely derived from low degree partial melting of a moderately thickened ancient lower crust. The diabases and the slightly postdated high Sr/Y granites in this study represent the magmatic responses to the final orogenic evolution in the southwestern XMOB. To-gether with regional works, we propose that the slab break-off of the Paleo-Asian oceanic lithosphere following the terminal collision between the North China Craton and the South Mongolia terranes triggered asthenospheric upwelling, and the ongoing convergence further initiated moderately crustal thickening and uplift in the XMOB.

0.   INTRODUCTION
  • Accretionary orogens formed at sites of subduction of the oceanic lithosphere are ultimately involved in a collisional phase when plate subduction ceased (Cawood et al., 2009), leading to continental amalgamation/collision, slab break-off, crustal thickening and subsequent lithospheric delamination, accompanied by extensive and diverse magmatism (e.g., Condie et al., 2009; Ferrari, 2004). Consequently, identification of magmatic records in response to these processes can help us better understand the final orogenic evolution in an accretionary orogen.

    The Central Asian orogenic belt (CAOB) (Fig. 1a) is one of the world's largest and best-preserved accretionary orogens (Wilde, 2015; Xiao et al., 2015, 2003; Windley et al., 2007; Jahn et al., 2000). The southeastern segment of the CAOB, which is also known as Xing-Meng orogenic belt (XMOB), was mainly constructed by the convergence between the South Mongolia terranes (SMT) to the north and the North China Craton (NCC) to the south, driven by the evolution of the Paleo-Asian Ocean from Neoproterozoic to Late Paleozoic (Zhao et al., 2018; Ma et al., 2017; Li et al., 2016a; Wilde, 2015; Xiao et al., 2015, 2003; Xu et al., 2013; Jian et al., 2010; Windley et al., 2007). The Solonker suture zone is widely considered to have marked the terminal closure of the Paleo-Asian Ocean during the end-Permian to Early-Middle Triassic (Li et al., 2017, 2016a; Xiao et al., 2015, 2003; Eizenhöfer et al., 2014; Schulmann and Paterson, 2011). However, unlike the well documented Paleozoic arc magmatism and oceanic subduction (e.g., Li et al., 2016b; Eizenhöfer et al., 2015, 2014; Xiao et al., 2015; Xu et al., 2013; Jian et al., 2010, 2008; Zhang et al., 2009a, b), origin and petrogenesis of the end-Permian to Early-Middle Triassic collision-related magmatism in the XMOB are still enigmatic. Recent studies propose that the end-Permian to Early Triassic magmatic rocks in Mandula and southeastern Xilinhot were generated in relation to the slab break-off following the terminal collision (Li et al., 2016a; Jian et al., 2010), and the Early-Middle Triassic high Sr/Y granitoids in Linxi and central Jilin were magmatic responses to the subsequent crustal thickening (Li et al., 2017; Wang et al., 2015a). However all those works were centered on the Solonker suture zone, the end-Permian to Early-Middle Triassic magmatism in southwestern XMOB receives much less attention.

    Figure 1.  (a) Simplified sketch map of the Central Asian orogenic belt showing major tectonic sub-divisions of Central Asia (modified after Jahn et al., 2000); (b) geological sketch map of the central-western Xing-Meng orogenic belt (modified after Xiao et al., 2015; Jian et al., 2008). Abbreviations: OSAC. Ondor Sum subduction-accretionary complex; BAAC. Baolidao arc-accretionary complex; HOAC. Hegenshan ophiolite-accretionary complex.

    In this study, we present new zircon U-Pb ages and Lu-Hf isotopes, whole rock geochemical and Sr-Nd-Pb isotopic data for the Early-Middle Triassic diabases and monzogranites in the Langshan area, southwestern XMOB (Fig. 1b). Our results, in combination with regional studies, enable us to propose that the studied diabases and the monzogranites were respectively generated in response to the slab break-off and the following crustal thickening during the final orogenic evolution in southwestern XMOB.

1.   GEOLOGICAL BACKGROUND AND SAMPLES
  • The XMOB has long been regarded as a complex collage of arcs, microcontinents, remnants of oceanic crust and associated volcanic-sedimentary sequences that amalgamated between the NCC and the SMT during the evolution and closure of the Paleo-Asian Ocean (e.g., Xiao et al., 2015, 2003; Jahn et al., 2000). The main part of the XMOB, from south to north, consists of the southern accretionary zone (SAZ) between the north margin of the NCC and the Solonker suture zone, the Solonker suture zone, the northern accretionary zone (NAZ) and the Hegenshan ophiolite-accretionary complex (Xiao et al., 2015; Jian et al., 2010). The SAZ includes the Early Paleozoic Bainaimiao arc (Zhang et al., 2014c) and the Ondor Sum subduction-accretion complex (Jian et al., 2008; de Jong et al., 2006; Xiao et al., 2003), the NAZ is characterized by the Paleozoic Baolidao arc-accretionary complex with accreted Precambrian blocks (e.g., Xilin Gol complex) (Xiao et al., 2015; Chen et al., 2009). The Hegenshan ophiolite-accretionary complex possesses a relatively complete lithological sequence (Miao et al., 2008), which is suggested to be formed in a subduction- related tectonic setting, such as back-arc or island arc-marginal basin, that related to the Early Carboniferous north-dipping subduction of the Paleo-Asian Ocean (Eizenhöfer et al., 2015; Miao et al., 2008; Robinson et al., 1999).

    The Solonker suture zone, which strikes from the Solonker area to the Linxi area (Fig. 1b) (Xiao et al., 2015; Jian et al., 2010, 2008), marks the final closure of the Paleo-Asian Ocean in the southeastern CAOB (e.g., Li et al., 2017, 2016a; Zhang et al., 2017; Eizenhöfer et al., 2015, 2014; Wilde, 2015; Xiao et al., 2015, 2003). The ages of the tectonic mélanges and ophiolitic fragments that discontinuously distributed along the Solonker suture zone are consistently Early Permian (Song et al., 2015; Jian et al., 2010). The Middle Permian Zhesi Formation volcaniclastic and turbiditic strata in the Solonker suture zone are unconformably overlain by the Late Permian-Early Triassic Linxi Formation clastic rocks, which are considered to be accumulated in a terrestrial setting (Eizenhöfer et al., 2015; Li et al., 2014). Zhang et al. (2016a) suggests that the Solonker suture zone underwent intermediate-to-low P/T type metamorphism during the Triassic. Several end-Permian to Early- Middle Triassic collision-related granitoids also have been identified in the Linxi area (Li et al., 2017, 2016a).

    The Langshan area in southwestern XMOB is approximately 100 km south of the Solonker suture zone (Fig. 1b). The Paleozoic tectonic affinity of the Langshan area was disputed because of the controversial relationship between the NCC and the Alxa Block (Zhang et al., 2016b; Yuan and Yang, 2015). Many recent studies have revealed that the Alxa Block had collided with the NCC before the Late Paleozoic (e.g., Wang et al., 2016; Zhang et al., 2016b), although the accurate timing and location of this amalgamation are still in debate. The eldest rocks in the Langshan area is the Archean high-grade metamorphic complex (Fig. 2), which are dominated by amphibolite, quartz schist and schistose to magmatic gneiss with granitic gneiss of 2 563 and 2 619 Ma (Liu, 2012). The Proterozoic basement rocks in the Langshan area mainly consist of greenschist-amphibolite facies metamorphosed carbonates, clastic rocks, with some Neoproterozoic meta-volcanic interlayers (Hu et al., 2014; Peng et al., 2010). Outcrops of the Paleozoic stratigraphy are only limited to the west of the Langshan area. They are mainly composed of some Carboniferous-Permian marine sedimentary and volcanic rocks, and are uncomfortably covered by the Jurassic-Quaternary sedimentary strata (Fig. 2). The Langshan area is also characterized by a complicated framework of Mesozoic multi-stage strike-slip faults and thrusts, which are interpreted to be related to the intraplate deformation phases across the southern CAOB (Zhang et al., 2016b, 2013; Darby and Ritts, 2007).

    Figure 2.  Geological sketch map of the Langshan area (modified after Liu et al., 2016).

    Igneous rocks in the Langshan area are mainly composed of the Late Paleozoic and Early Mesozoic felsic intrusions (Fig. 2) (e.g., Liu et al., 2016; Wang et al., 2015b; Lin et al., 2014; Peng et al., 2013), such as the ~338 Ma Baoribu quartz diorite with εHf(t) values of -6.9 to +2.0 in the north Chaoge Ondor (Liu et al., 2016), the ~281 Ma deformed granitic porphyry in the Urad Houqi area (Lin et al., 2014), and the ~259 Ma granodiorite with εNd(t) values of -2.7 ~ -1.6 close to the Huogeqi ore district (Wang et al., 2015b). Additionally, some small outcrops of the Early Paleozoic granitoids and the Late Paleozoic mafic intrusions are also observed (e.g., Liu et al., 2016; Wang et al., 2016, 2015b).

    The investigated diabases are located immediately to the east of the Chaoge Ondor (Fig. 2), most of them are vertical and NE-SW trending. They intruded into the Late Paleozoic granitic rocks (Fig. 3a) and the Proterozoic greenschist-amphibolite facies metamorphic clastic rocks (Zhaertai Group), but did not penetrate the Mesozoic granitoids in the field. The intrusive contacts are very clear and thermal recrystallization can be observed near the boundary of most dykes. In contrast to their country rocks, the diabases are rarely subjected to metamorphism and deformation. Individual diabase dykes in the Langshan area typically have widths of 0.8 to 3.0 m and can be traced for distances of 0.2 to 8 km. The investigated monzogranites are located approximately 15 km west of Chaoge Ondor, with surface outcrops over 60 km2 (Figs. 2 and 3b). They mainly intruded into the Early Carboniferous quartz diorite in the field, and are devoid of any effects of metamorphism or deformation. The diabases display ophitic or sub-ophitic textures, with similar mineral compositions of lath-shaped plagioclase (50 vol.%-60 vol.%) and pyroxene (30 vol.%-40 vol.%), with minor hornblende (~5 vol.%) and Fe-Ti oxides (Fig. 3c). The monzogranite is composed of plagioclase (25 vol.%-30 vol.%), K-feldspar (30 vol.%-35 vol.%), quartz (25 vol.%-30 vol.%), biotite (3 vol.%-10 vol.%) and minor accessory minerals (e.g., magnetite, zircon and apatite) (Fig. 3d).

    Figure 3.  Representative field photographs and micrographs of the studied diabase and monzogranite from the Langshan area. Mineral abbreviations: Py. pyroxene; Hbl. hornblende; Pl. plagioclase; Kf. K-feldspar; Bt. biotite; Qtz. quartz.

2.   ANALYTICAL METHODS
  • Zircon grains were separated from the whole-rock samples by standard techniques of density and magnetic methods and further purified by hand-picking using a binocular microscope. The internal structures of each grain were revealed by Cathodoluminescence (CL) imaging prior to the U-Pb analysis. Zircon U-Pb analysis for sample Ls23-d2-tw2 was carried out at the State Key Laboratory of Continental Dynamics, Northwest University, Xi'an, China, by using an Agilent 7500a ICP-MS equipped with a 193 nm laser, the spot size is 35 μm. Zircon U-Pb age analysis for sample Ls06-d1-tw1 was carried out on a Finnigan Neptune ICP-MS that equipped with New Wave 193 nm excimer lasers, at the Tianjin Institute of Geology and Mineral Resources, Tianjin, China. A 35 μm diameter spot was used for the laser ablation of a single zircon grain. The 206Pb/238U and 207Pb/206Pb ratios were calculated by the GLITTER program and corrections were applied using Harvard zircon 91500 as an external calibration. Details of the analytical methodology are described by Li et al. (2009) and Yuan et al. (2004). The common Pb corrections were made following the method of Andersen (2002). Isoplot 3.0 program (Ludwig, 2003) was used for data processing.

    Whole-rock geochemical compositions were determined at the Analytical Laboratory of the Beijing Research Institute of Uranium Geology (BRIUG), China. Major elements were analysed on fused glass discs by a Philips PW2404 X-ray fluorescence spectrometry with analytical errors better than 1%-5%, FeO contents were determined using the classical wet chemical method. Trace elements were determined by using a Perkin- Elmer Sciex ELAN DRC-e inductively coupled plasma mass spectrometer (ICP-MS) and the analytical uncertainties were mostly better than 5%. Analytical procedures and instrument parameters were similar to those described by Gao et al. (2002).

    The whole rock Sr-Nd and Pb isotopic analysis were performed on triton thermal ionization mass spectrometer (TIMS) and MAT261 TIMS, respectively, at the Central-South China Supervision and Inspection Center of Mineral Resources, Ministry of Land and Resources, Wuhan, China. The mass fractionation of Sr isotopic ratios has been exponentially corrected to the 86Sr/88Sr=0.119 4 and Nd isotopic ratios to 146Nd/144Nd= 0.721 9. The values for the NBS 987 Sr-standard and the BCR-2 Nd-standard measured during the data acuisition were 87Sr/86Sr=0.710 31±4 (2σ) and 143Nd/144Nd=0.512 62±5 (2σ), respectively. Repeated analysis yielded the 208Pb/206Pb, 207Pb/206Pb and 204Pb/206Pb ratios of the Standard NBS981 in this study were 2.164 91±13 (2σ), 0.914 37±5 (2σ) and 0.059 104±5 (2σ), respectively. Detailed descriptions of the analytical techniques are given by Zeng et al. (2013).

    In-situ zircon Lu-Hf isotopic analysis was conducted by using a Nu Plasma Ⅱ MC-ICP-MS coupled with a Resolution M-50 193 nm laser at the State Key Laboratory of Continental Dynamics, Northwest University, Xi'an, China. A 44 μm laser beam with a repetition rate of 6 Hz was used for the analysis. Detailed information of these instruments and analytical technique can be found in Bao et al. (2017) and Yuan et al. (2008). The 91500 and MT standards yielded weighted 176Hf/177Hf ratios of 0.282 264±0.000 030 and 0.282 528±0.000 019, respectively.

3.   RESULTS
  • Zircons selected from diabase sample Ls23-d2-tw2 and monzogranite sample Ls06-d1-tw1 are mostly subhedral to euhedral, colourless or buff to transparent, with lengths of 50-200 μm. The majority of zircon grains exhibit clearly oscillatory zoning in the CL images (Fig. 4), suggesting their magmatic origin (Hoskin and Schaltegger, 2003). The Th/U ratios of zircons from Ls23-d2-tw2 and Ls06-d1-tw1 range from 0.04-1.34 and 0.06-0.42, respectively. Detailed zircon U-Pb isotopic data are presented in Table 1. Twenty-one spots were conducted on zircons from sample Ls23-d2-tw2, and most of them are located below the concordia line (Fig. 5a). However, four relatively younger zircon grains have concordant ages that range from 248 to 255 Ma and yield a weighted mean 206Pb/238U age of 251.4±3.6 Ma, and other five grains have concordant ages ranging from 271 to 273 Ma, yielding a concordant age of 272.1±1.6 Ma (Fig. 5b). A total of twenty-five spots of zircons from the monzogranite sample Ls06-d1-tw1 show nearly concordant ages and yield a well-weighted mean 206Pb/238U age of 245.5±0.7 Ma (Figs. 5c-5d), which means that emplacement of the studied monzogranite took place in the Middle Triassic.

    Figure 4.  (a), (b) CL images of representative zircons from the studied diabases and (c) monzogranites. The solid and dotted circles denote the spots for U-Pb dating and Lu-Hf, respectively.

    No. Isotopic ratios Ages (Ma)
    Th/U 206Pb/238U 1σ 207Pb/235U 1σ 207Pb/206Pb 1σ 206Pb/238U 1σ 207Pb/235U 1σ 207Pb/206Pb 1σ
    Ls23-d2-tw2 (N41°25′46″/E106°54′02″)
    1 0.68 0.039 80 0.000 50 0.279 70 0.013 40 0.050 90 0.002 50 252 3 250 11 238 117
    2 1.26 0.043 00 0.000 80 0.308 50 0.020 20 0.052 10 0.003 60 271 5 273 16 288 117
    3 0.91 0.040 30 0.000 60 0.277 20 0.012 10 0.049 90 0.002 40 255 4 248 10 188 74
    4 0.73 0.039 60 0.000 60 0.282 90 0.021 80 0.051 90 0.004 10 250 4 253 17 280 180
    5 0.29 0.328 60 0.003 70 5.492 70 0.067 90 0.121 20 0.002 60 1 832 18 1 899 11 1 974 10
    6 0.71 0.065 90 0.000 90 0.528 10 0.018 50 0.058 20 0.002 30 411 6 431 12 536 52
    7 0.45 0.043 20 0.000 60 0.315 60 0.010 30 0.053 00 0.002 00 273 3 279 8 330 50
    8 0.68 0.039 20 0.000 60 0.296 00 0.015 20 0.054 70 0.003 00 248 4 263 12 401 86
    9 1.34 0.122 70 0.001 70 1.140 80 0.033 80 0.067 50 0.002 30 746 10 773 16 852 39
    10 0.04 0.090 90 0.001 00 0.938 50 0.014 90 0.074 90 0.001 50 561 6 672 8 1 065 40
    11 0.68 0.043 30 0.000 60 0.355 60 0.010 80 0.059 60 0.002 10 273 3 309 8 591 43
    12 0.6 0.067 80 0.000 80 0.524 50 0.010 80 0.056 10 0.001 50 423 5 428 7 456 26
    13 0.72 0.071 20 0.000 80 0.592 80 0.010 40 0.060 40 0.001 50 444 5 473 7 617 20
    14 1.08 0.197 50 0.002 20 2.696 00 0.035 20 0.099 00 0.002 20 1 162 12 1 327 10 1 606 11
    15 0.63 0.212 00 0.002 80 2.646 30 0.081 30 0.090 50 0.003 00 1 240 15 1 314 23 1 437 65
    16 0.51 0.043 00 0.000 70 0.334 10 0.016 80 0.056 40 0.003 10 271 4 293 13 467 83
    17 0.14 0.095 40 0.001 10 1.015 80 0.019 80 0.077 20 0.002 00 588 7 712 10 1 127 21
    18 0.71 0.043 00 0.000 60 0.326 80 0.015 20 0.055 10 0.002 70 271 4 287 12 416 111
    19 0.44 0.186 10 0.002 20 2.095 90 0.046 20 0.081 70 0.002 00 1100 12 1 147 15 1 238 50
    20 0.68 0.337 40 0.004 30 5.418 60 0.097 20 0.116 50 0.002 90 1874 21 1 888 15 1 903 16
    Ls06-d1-tw1 (N41°26′16″/E106°44′30″)
    1 0.91 0.038 77 0.000 36 0.280 13 0.009 80 0.052 41 0.001 88 245 2 251 9 303 82
    2 0.73 0.038 82 0.000 36 0.273 59 0.005 48 0.051 12 0.000 89 246 2 246 5 246 40
    3 0.29 0.038 93 0.000 31 0.276 62 0.008 70 0.051 53 0.001 62 246 2 248 8 265 72
    4 0.71 0.038 78 0.000 40 0.273 56 0.019 43 0.051 16 0.003 77 245 3 246 17 248 170
    5 0.45 0.038 84 0.000 37 0.275 39 0.020 03 0.051 42 0.003 68 246 2 247 18 260 164
    6 0.68 0.039 14 0.000 30 0.277 27 0.005 75 0.051 38 0.001 00 248 2 248 5 258 45
    7 1.34 0.038 43 0.000 23 0.454 56 0.009 99 0.085 79 0.001 80 243 1 380 8 1 333 41
    8 0.04 0.039 42 0.000 24 0.275 38 0.005 73 0.050 66 0.000 99 249 2 247 5 226 45
    9 0.68 0.038 24 0.000 28 0.280 37 0.009 66 0.053 17 0.001 72 242 2 251 9 336 73
    10 0.6 0.038 95 0.000 23 0.279 60 0.006 37 0.052 07 0.001 17 246 1 250 6 288 51
    11 0.72 0.038 98 0.000 23 0.279 26 0.006 01 0.051 96 0.001 10 246 1 250 5 284 49
    12 1.08 0.038 76 0.000 24 0.278 24 0.005 40 0.052 07 0.000 99 245 2 249 5 288 43
    13 0.63 0.039 05 0.000 24 0.280 64 0.007 87 0.052 12 0.001 43 247 2 251 7 291 63
    14 0.51 0.038 82 0.000 25 0.274 83 0.012 22 0.051 34 0.002 25 246 2 247 11 256 101
    15 0.14 0.038 92 0.000 22 0.275 04 0.004 93 0.051 25 0.000 92 246 1 247 4 252 41
    16 0.71 0.038 66 0.000 24 0.273 00 0.014 41 0.051 21 0.002 68 245 2 245 13 250 121
    17 0.44 0.038 96 0.000 22 0.282 89 0.003 31 0.052 66 0.000 60 246 1 253 3 314 26
    18 0.35 0.038 87 0.000 23 0.273 73 0.006 22 0.051 08 0.001 13 246 1 246 6 244 51
    19 0.32 0.038 91 0.000 26 0.278 55 0.016 33 0.051 92 0.003 02 246 2 250 15 282 133
    20 0.28 0.038 90 0.000 22 0.280 12 0.006 36 0.052 23 0.001 16 246 1 251 6 295 51
    21 0.22 0.038 34 0.000 22 0.282 48 0.006 86 0.053 44 0.001 27 243 1 253 6 348 54
    22 0.06 0.038 69 0.000 25 0.275 32 0.010 41 0.051 62 0.001 93 245 2 247 9 268 86
    23 0.31 0.038 34 0.000 22 0.413 30 0.012 61 0.078 18 0.002 32 243 1 351 11 1 151 59
    24 0.4 0.038 93 0.000 26 0.273 36 0.016 08 0.050 93 0.002 95 246 2 245 14 237 133
    25 0.42 0.039 02 0.000 23 0.279 43 0.007 54 0.051 94 0.001 39 247 1 250 7 283 61

    Table 1.  LA-ICP-MS U-Pb isotopic data for the zircons from the studied diabase and monzogranite samples

    Figure 5.  (a), (b) U-Pb concordia diagrams of zircon grains from the studied diabase sample Ls23-d2-tw2; (c) (d) monzogranite sample Ls06-d1-tw1. (b) shows the results of zircons with 206Pb/238U ages less than 600 Ma for Ls23-d2-tw2.

  • The major and trace element concentrations of studied diabase and monzogranite samples are listed in Table 2. The relatively low LOI (loss on ignition) values of most samples indicate a limited degree of hydrothermal alteration.

    Sample Ls01-d1-q2 Ls01-d1-q10 Ls01-d1-q17 Ls01-d2-q7 Ls01-d2-q10 Ls01-d3-q5 Ls01-d1-q8 Ls01-d1-q14 Ls03-d6-q1 Ls06-d1-q1 Ls06-d3-q1 PM007D4-1GS PM007D4-1W
    Rock type Diabase Diabase Diabase Diabase Diabase Diabase Diabase Diabase Diabase Monzogranite Monzogranite Monzogranite Monzogranite
    Location N41°25′29″ E106°54′15″ N41°25′32″ E106°54′10″ N41°25′39″ E106°54′14″ N41°25′41″ E106°53′52″ N41°25′41″ E106°53′50″ N41°26′01″ E106°53′42″ N41°25′31″ E106°54′11″ N41°25′35″ E106°54′07″ N41°26′18″ E106°54′36″ N41°26′16″ E106°44′30″ N41°25′56″ E106°44′01″ N41°25′32″ E106°44′25″ N41°25′32″ E106°44′25″
    Major elements (wt.%)
    SiO2 52.14 52.38 54.55 51.45 51.13 52.92 53.21 53.13 52.22 70.09 70.27 72.55 72.63
    TiO2 1.35 1.14 1.32 1.38 1.27 1.40 1.45 1.51 1.86 0.32 0.33 0.25 0.26
    Al2O3 17.41 15.01 15.53 18.20 17.03 17.06 20.83 21.57 18.42 15.07 14.79 14.17 14.10
    Fe2O3 7.90 2.61 2.53 2.89 2.39 2.51 2.06 2.40 8.18 1.30 1.39 0.99 1.00
    FeO 1.00 5.68 5.15 5.42 5.79 5.57 4.60 3.15 1.06 1.08 1.02 0.65 0.70
    MnO 0.12 0.12 0.11 0.12 0.12 0.13 0.09 0.07 0.12 0.05 0.05 0.02 0.03
    MgO 5.75 9.66 7.53 5.58 6.66 5.26 3.01 2.27 3.78 0.86 0.87 0.40 0.42
    CaO 8.00 7.81 6.63 8.88 8.72 7.62 7.91 7.65 7.02 2.26 2.28 1.37 1.40
    Na2O 3.82 2.56 3.34 3.36 3.55 4.09 4.11 4.07 4.23 4.08 3.98 4.07 4.06
    K2O 1.20 1.31 1.06 0.94 0.76 1.54 1.84 1.75 1.58 3.76 3.96 4.67 4.58
    P2O5 0.18 0.17 0.18 0.18 0.16 0.20 0.24 0.24 0.31 0.10 0.11 0.07 0.08
    LOI 1.82 1.77 2.27 1.83 2.55 1.91 0.84 2.34 2.01 1.38 1.22 0.81 0.77
    Trace elements (ppm)
    Sc 22.1 22.0 20.7 25.9 25.7 27.0 15.1 15.2 19.8 4.1 4.6 2.7 2.8
    V 128.0 145.0 132.0 121.0 149.0 161.0 102.0 100.0 128.0 23.0 31.5 36.3 37.7
    Cr 59.2 494.0 268.0 82.4 158.0 139.0 62.4 53.6 66.3 4.7 8.3 5.7 5.8
    Co 29.7 33.2 25.1 31.3 32.9 24.3 21.1 16.0 25.8 3.1 4.1 2.1 2.1
    Ni 30.8 123.0 88.0 26.9 57.6 31.3 30.6 24.5 36.2 3.6 6.4 1.2 1.9
    Cu 30.7 29.4 50.8 40.3 39.5 45.8 49.4 51.2 64.4 7.5 7.8 12.3 13.2
    Zn 70.5 66.6 63.4 77.2 75.3 74.0 69.4 61.9 89.1 32.9 41.6 47.6 49.9
    Ga 16.8 14.5 15.8 19.2 18.5 18.1 20.7 19.7 20.7 13.8 16.2 21.3 21.4
    Cs 4.3 4.0 4.2 2.8 2.2 1.4 5.7 5.9 2.5 5.5 4.9 4.9 5.0
    La 11.8 11.3 14.5 12.0 11.5 12.4 17.9 16.7 21.4 19.7 20.1 34.3 39.7
    Ce 26.2 24.8 30.8 26.7 26.0 28.3 37.0 36.0 46.6 35.1 37.6 58.2 67.5
    Pr 3.62 3.44 4.12 3.75 3.50 3.92 5.24 5.02 6.21 3.81 3.89 5.97 7.01
    Nd 17.2 16.0 18.3 17.4 16.0 18.2 23.5 22.7 28.1 13.7 13.9 20.0 22.2
    Sm 3.99 3.76 4.13 4.47 4.05 4.71 5.41 5.44 6.71 2.43 2.21 2.97 3.24
    Eu 1.50 1.24 1.33 1.57 1.44 1.47 1.90 1.85 1.83 0.49 0.54 0.72 0.72
    Trace elements (ppm)
    Gd 3.88 3.47 3.69 4.11 3.76 4.23 5.18 4.98 6.04 2.07 1.93 2.14 2.40
    Tb 0.82 0.69 0.79 0.83 0.78 0.87 1.03 1.06 1.25 0.36 0.31 0.30 0.35
    Dy 4.63 3.91 4.31 4.92 4.54 5.10 5.79 5.96 6.79 1.84 1.48 1.49 1.80
    Ho 0.93 0.79 0.85 0.97 0.91 1.02 1.16 1.19 1.34 0.35 0.29 0.25 0.26
    Er 2.60 2.26 2.49 2.77 2.55 2.86 3.27 3.37 3.78 1.07 0.85 0.68 0.76
    Tm 0.44 0.37 0.40 0.45 0.44 0.46 0.52 0.55 0.60 0.16 0.15 0.10 0.10
    Yb 2.72 2.31 2.51 2.81 2.72 2.97 3.22 3.34 3.87 1.15 1.01 0.60 0.70
    Lu 0.39 0.33 0.35 0.38 0.36 0.40 0.45 0.46 0.50 0.17 0.16 0.09 0.10
    Rb 35.5 48.1 44.1 40.4 29.9 73.2 83.2 71.7 72.2 109.0 128.0 169.0 173.0
    Sr 339 287 343 413 352 405 438 391 398 310 383 320 324
    Y 23.10 21.00 22.20 27.40 25.00 27.80 32.40 32.80 37.50 10.50 9.26 7.16 8.03
    Zr 159 148 160 179 165 181 212 209 285 116 142 120 142
    Nb 4.24 4.10 5.61 4.99 4.40 4.65 5.77 6.10 7.01 8.44 9.22 9.13 9.60
    Ba 233 131 163 133 116 223 229 195 255 658 991 785 770
    Hf 4.39 3.93 4.33 4.53 4.32 4.92 5.57 5.43 7.22 3.73 3.88 3.13 3.86
    Ta 0.39 0.36 0.49 0.38 0.37 0.36 0.49 0.51 0.44 0.87 0.87 1.07 1.13
    Pb 5.69 5.61 7.79 5.47 6.44 6.16 9.02 9.72 9.56 22.00 27.50 31.70 31.70
    Th 2.77 2.41 3.48 2.62 2.74 2.98 4.27 4.28 5.60 8.69 8.87 22.20 25.60
    U 0.74 0.62 0.92 1.24 0.78 0.83 1.90 1.27 1.75 1.92 1.50 1.45 1.76

    Table 2.  Major oxide and trace elements composition of the studied diabases and monzogranites

    The diabase samples have SiO2 contents of 51.13 wt.%- 54.55 wt.%, K2O contents of 0.76 wt.%-1.84 wt.% and Na2O contents of 2.56 wt.%-4.23 wt.%. They plot in the compositional fields of gabbro, gabbroic diorite, monzogabbro and monzodiorite (Fig. 6a) in the TAS diagram, and belong to the calc-alkaline series (Figs. 6b-6d). These characteristics are similar to those of the ~250 Ma E-MOEB-like diabase from the Solonker ophiolitic mélanges (Jian et al., 2010). In the Harker diagrams, they exhibit regular trends of decreasing Al2O3, Na2O, K2O, TiO2 and P2O5 contents with increasing MgO contents (Fig. 7). All the studied diabase samples show similar rare earth element (REE) compositions, with total REE contents ranging from 74.7 ppm to 135.0 ppm. In the chondrite-normalized REE diagram (Fig. 8a), they are characterized by light REE (LREE) enrichment and slightly negative to positive Eu anomalies, with (La/Yb)N of 2.99-4.14 and Eu/Eu* of 0.88-1.17. Their primitive mantle-normalized trace element patterns (Fig. 8b) are generally characterized by enrichment in large ion lithophile elements (LILE; e.g., Rb, U, K, Pb) and depletion in high field strength elements (HFSE; e.g., Nb, Ta, P, Ti), which are similar to those of the "arc-like" magmatic rocks from the subduction zone (Ma et al., 2014).

    Figure 6.  (a) TAS diagram (Middlemost, 1994; Irvine and Baragar, 1971); (b) K2O vs. SiO2 diagram (Rickwood, 1989); (c) A/NK vs. A/CNK diagram (Peccerillo and Taylor, 1976); (d) AFM diagram (Irvine and Baragar, 1971). The previously reported ~250 Ma Solonker E-MORB like diabases (Jian et al., 2010), the Linxi Triassic high-Mg adakitic andesites (Liu et al., 2012) and the Early-Middle Triassic High Sr/Y granitoids (Li et al., 2017) are plotted for comparison.

    Figure 7.  Harker diagrams of MgO vs. selected major and trace elements for the studied diabase samples.

    Figure 8.  (a) Chondrite-normalized REE diagram and (b) primitive mantle- normalized trace element diagram for the studied diabase and monzogranite samples. The values of chondrite and primitive mantle are from Boynton (1984) and Sun and McDonough (1989), respectively. Data for the Early- Middle Triassic high Sr/Y granitoids in Linxi area are from Li et al. (2017).

    The monzogranite samples exhibit relatively high SiO2 (70.09 wt.%-72.63 wt.%), Na2O (3.98 wt.%-4.08 wt.%), K2O (3.76 wt.%-4.67 wt.%) and low Fe2O3T (1.54 wt.%-2.27 wt.%), MgO (0.40 wt.%-0.87 wt.%), TiO2 (0.25 wt.%-0.33 wt.%) contents and Mg# (100×MgO/(MgO+TFeO)) values (31.6- 40.5). These rocks belong to high-K calc-alkaline series, and are metaluminous to weakly peraluminous with A/CNK ratios of 0.99-1.01. They also have relatively high concentrations of Al2O3 (14.10 wt.%-15.07 wt.%), Sr (310 ppm-380 ppm) but low Y (7.16 ppm-10.50 ppm) and Yb (0.60 ppm-1.15 ppm) contents, with high Sr/Y ratios (30-45) (Fig. 9). The monzogranites show notable enrichment in LREE ((La/Yb)N=12.29- 40.80) and slightly negative Eu anomalies (Eu/Eu*=0.66-0.88) in the chondrite-normalized diagram (Fig. 8a). Their primitive mantle-normalized trace element diagrams are generally characterized by enrichment in LILE, such as Rb, U and K, and depletion in Nb, Ta, Ti, and P (Fig. 8b). Overall, the monzogranites in this study display geochemical characteristics similar to those of the coeval adakitic rocks along the Solonker suture zone, such as the Early-Middle Triassic high Sr/Y granitoids in the Linxi area (Li et al., 2017) and the Early Triassic adakitic granitoids in the central Jilin area (Wang et al., 2015a).

    Figure 9.  Adakite discrimination diagrams (a) Sr/Y vs. Y and (b) (La/Yb)N vs. YbN for the studied monzogranite samples. Fields for adakite and arc calc-alkaline rocks are based on Petford and Atherton (1996) and Defant and Drummond (1990). The crystal fractionation paths of the primary minerals in Fig. 9a are from Castillo et al. (1999). The batch partial melting trends in Fig. 9b are based on Petford and Atherton (1996).

  • Detailed results of whole-rock Rb-Sr, Sm-Nd and Pb isotopes for the studied diabases are given in Table 3. Initial values of all the samples have been calculated based on their youngest concordant zircon U-Pb age of ~251 Ma. As shown in Fig. 10, the diabase samples have relatively variable initial 87Sr/86Sr values of 0.705 86 to 0.708 87 and εNd(t) values of -3.1 to +1.5, with Nd model ages ranging from 1 473 to 1 034 Ma, these characteristics are apparently different from the end-Permain mafic rocks in the Solonker area (Luo et al., 2016) and the Early Mesozoic ultramafic-mafic complexes along the northern NCC (Zhang et al., 2012), but similar to those of the Late Paleozoic enriched mantle-derived mafic intrusions in the region (Wang et al., 2015b). As for the Pb isotopic compositions, the initial ratios of 208Pb/204Pb, 207Pb/204Pb and 206Pb/204Pb of the diabase samples vary from 35.968-37.346, 15.448-15.508 and 16.280-17.492, respectively, they plot above the North Hemisphere Reference Line (NHRL) and close to the compositional field of EM1 (Fig. 11).

    Sample Ls01-d1-2 Ls01-d1-8 Ls01-d1-17 Ls01-d2-7 Ls03-d6-1
    Rb (μg/g) 39.46 86.87 48.36 39.51 75.61
    Sr (μg/g) 376.9 452.3 385 412.1 414.1
    87Rb/86Sr 0.301 8 0.553 9 0.362 2 0.276 3 0.526 5
    87Sr/86Sr 0.707 56 0.709 84 0.710 16 0.706 85 0.710 26
    1δ 0.000 003 0.000 003 0.000 005 0.000 003 0.000 003
    (87Sr/86Sr)i 0.706 483 0.707 863 0.708 867 0.705 864 0.708 381
    Sm (μg/g) 5.13 5.474 4.319 4.163 6.707
    Nd (μg/g) 20.3 23.24 18.58 16.6 28.42
    147Sm/144Nd 0.152 9 0.142 5 0.140 6 0.151 7 0.142 8
    143Nd/144Nd 0.512 668 0.512 556 0.512 48 0.512 717 0.512 618
    1δ 0.000 005 0.000 005 0.000 004 0.000 004 0.000 004
    (143Nd/144Nd)i 0.512 421 0.512 326 0.512 253 0.512 472 0.512 388
    εNd(t) 2.06 0.19 -1.23 3.05 1.4
    TDM (Ma) 1 076 1 158 1 283 1 473 1 034
    Pb (μg/g) 5.47 1.54 3 1.38 3.12
    U (μg/g) 0.97 0.58 0.8 0.55 1.55
    Th (μg/g) 2.89 2.17 2.79 2.21 5.09
    206Pb/204Pb 18.508 18.524 18.48 18.591 18.587
    207Pb/204Pb 15.55 15.533 15.52 15.523 15.545
    208Pb/204Pb 38.523 38.549 38.438 38.429 38.603
    (206Pb/204Pb)i 17.492 16.841 17.363 16.544 16.28
    (207Pb/204Pb)i 15.508 15.453 15.476 15.451 15.448
    (208Pb/204Pb)i 37.346 36.042 36.825 35.628 35.968

    Table 3.  Whole rocks Sr-Nd-Pb isotopic data for the studied diabase samples

    Figure 10.  Plot of εNd(t) vs. (87Sr/86Sr)i for the studied diabase samples. The fields for DM (depleted mantle), MORB (middle ocean ridge basalt) and OIB (ocean island basalt) are from Zindler and Hart (1986). Data for the Late Paleozoic ultramafic-mafic complexes from the northern NCC are from Zhang et al. (2009b). Data for the Late Paleozoic enriched mantle-derived mafic intrusions in Langshan area and the end-Permian mafic rocks in the Solonker area are from Wang et al. (2015b) and Luo et al. (2016), respectively.

    Figure 11.  (a) (206Pb/204Pb)i vs. (207Pb/204Pb)i and (b) (206Pb/204Pb)i vs. (208Pb/204Pb)i diagrams for the studied diabase samples. The fields for I-MORB (Indian MORB), P & N-MORB (Pacific and North Atlantic MORB), OIB, EM1, NHLR and the 4.5 Ga geochron are from Zou et al. (2000), Zindler and Hart (1986) and Hart (1984).

  • The zircon grains of monzogranite sample Ls06-d1-tw1 which were previously used for U-Pb age dating were also selected for Lu-Hf isotopes on the same spot. Individual Lu-Hf isotopic data, as well as the 206Pb/238U ages of the corresponding spots, are given in Table 4. Twenty analyses on zircons from sample Ls06-d1-tw1 yielded initial 176Hf/177Hf ratios of 0.281 896 to 0.282 373, their εHf(t) values and the corresponding TDM2 range from -25.8 to -8.8 (Fig. 12) and 2 900 to 1 833 Ma, respectively.

    Sample spot Age (Ma) 176Hf/177Hf 2σ 176Lu/177Hf 2σ 176Yb/177Hf 2σ εHf(t) TDM1 (Ma) TDM2 (Ma) fLu/Hf
    Ls06-d1-tw1.01 245 0.282 271 0.000 024 0.020 840 0.000 081 0.000 859 0.000 003 -12.5 1 379 2 064 -0.97
    Ls06-d1-tw1.02 246 0.282 227 0.000 019 0.024 826 0.000 150 0.000 958 0.000 006 -14.0 1 444 2 163 -0.97
    Ls06-d1-tw1.03 246 0.282 351 0.000 019 0.015 781 0.000 135 0.000 623 0.000 005 -9.6 1 260 1 883 -0.98
    Ls06-d1-tw1.04 246 0.282 373 0.000 017 0.013 352 0.000 114 0.000 550 0.000 005 -8.8 1 227 1 833 -0.98
    Ls06-d1-tw1.05 248 0.281 896 0.000 020 0.034 959 0.000 328 0.001 350 0.000 012 -25.8 1 923 2 900 -0.96
    Ls06-d1-tw1.06 243 0.282 267 0.000 018 0.015 929 0.000 048 0.000 665 0.000 002 -12.6 1 378 2 073 -0.98
    Ls06-d1-tw1.07 242 0.282 247 0.000 020 0.017 003 0.000 428 0.000 685 0.000 016 -13.4 1 406 2 118 -0.98
    Ls06-d1-tw1.08 246 0.281 930 0.000 020 0.036 772 0.000 395 0.001 405 0.000 014 -24.6 1 879 2 827 -0.96
    Ls06-d1-tw1.09 246 0.282 232 0.000 016 0.018 090 0.000 126 0.000 729 0.000 005 -13.8 1 428 2 148 -0.98
    Ls06-d1-tw1.10 245 0.282 229 0.000 017 0.017 079 0.000 140 0.000 683 0.000 005 -13.9 1 431 2 156 -0.98
    Ls06-d1-tw1.11 247 0.282 232 0.000 020 0.016 948 0.000 131 0.000 691 0.000 005 -13.8 1 427 2 148 -0.98
    Ls06-d1-tw1.12 246 0.282 228 0.000 020 0.015 719 0.000 280 0.000 638 0.000 011 -13.9 1 430 2 157 -0.98
    Ls06-d1-tw1.13 245 0.282 306 0.000 020 0.013 909 0.000 099 0.000 565 0.000 004 -11.2 1 320 1 983 -0.98
    Ls06-d1-tw1.14 246 0.282 100 0.000 019 0.021 041 0.000 066 0.000 838 0.000 003 -18.5 1 616 2 445 -0.97
    Ls06-d1-tw1.15 246 0.282 351 0.000 018 0.013 326 0.000 214 0.000 527 0.000 008 -9.6 1 257 1 883 -0.98
    Ls06-d1-tw1.16 246 0.282 275 0.000 022 0.014 781 0.000 150 0.000 585 0.000 005 -12.3 1 364 2 053 -0.98
    Ls06-d1-tw1.17 246 0.282 155 0.000 021 0.018 748 0.000 121 0.000 748 0.000 004 -16.5 1 535 2 320 -0.98
    Ls06-d1-tw1.18 243 0.282 175 0.000 019 0.016 144 0.000 044 0.000 668 0.000 002 -15.9 1 504 2 277 -0.98
    Ls06-d1-tw1.19 245 0.282 040 0.000 019 0.024 647 0.000 176 0.000 986 0.000 007 -20.7 1 704 2 578 -0.97
    Ls06-d1-tw1.20 246 0.282 239 0.000 019 0.016 306 0.000 111 0.000 650 0.000 004 -13.6 1 416 2 134 -0.98

    Table 4.  Zircon Lu-Hf isotopic compositions of the studied monzogranite sample Ls06-d1-tw1

    Figure 12.  Zircon εHf(t) vs. U-Pb age diagram for the studied monzogranite sample Ls06-d1-tw1. Data source for the Triassic high Sr/Y granitoids in northern and southern Linxi area are from Li et al. (2017). The end-Permian gabbros in Solonker (Luo et al., 2016) and the Late Paleozoic intrusions in SW XMOB (Liu et al., 2016; Peng et al., 2013; Pi et al., 2010) are also shown for comparison.

4.   DISCUSSION
  • Since the diabase dykes in this study are mostly hosted by the Late Paleozoic granitoids and unconformably covered by the Early Cretaceous Guyang Formation in the field, their formation ages can be roughly restricted to Late Paleozoic to Early Mesozoic. Besides, the pre-Mesozoic geological units in the Langshan area are pervasively deformed (Zhang et al., 2013; Darby and Ritts, 2007), whereas the studied diabases are undeformed and unmetamorphosed. Thus, the timing of the youngest deformation in the region can provide constraint on the upper age limit of the studied diabase. As reported by Lin et al. (2014), the Early Permian (~290 Ma) granitic porphyries in Urad Houqi are generally characterized by NE-striking and N-vergent orientations and structures, and they are the youngest deformed intrusions in the Langshan area, which implies that the diabase dykes in this study must have been formed after Early Permian.

    Although zircons are volumetrically minor in mafic- ultramafic magma and zircon grains obtained from the mafic- ultramafic rocks are mostly captured, previous studies have shown that high-precision zircon U-Pb dating remains useful in placing constraints on the formation ages of mafic-ultramafic rocks (e.g., Yu et al., 2017; Zhang et al., 2009b). Our U-Pb dating results indicate that the zircon 206Pb/238U ages acquired from diabase sample Ls23-d2-tw2 are varying from 248 to 1 874 Ma. The older ages may be xenocrysts that potentially captured during the magma ascent, but nine relatively younger zircons yielded two weighted mean 206Pb/238U ages of ~272 and ~251 Ma (Fig. 5b). Considering that the Early Permian igneous rocks are widely distributed in the southwestern XMOB (e.g., Wang et al., 2015b; Lin et al., 2014; Peng et al., 2013; Zhang et al., 2009a), we propose that the weighted mean age of ~272 Ma was more likely to be in relation to the Early Permian magmatism in the region. The four zircons with ages of ~251 Ma are relatively larger in size and exhibit developed oscillatory zoning in the CL images (Fig. 4), these features are more akin to those of zircons from the intermediate-felsic magmatic rocks instead of the mafic magma, suggesting that the studied diabase probably formed after ~251 Ma. Moreover, previous studies show that extensive Middle-Late Triassic magmatism (including the Middle Triassic high Sr/Y monzogranite in this study) was developed in southwestern XMOB (e.g., Zhang et al., 2014a and references therein). However, no Middle-Late Triassic zircon is observed in sample Ls23-d2-tw2, which means that the studied diabase is unlikely to be formed after Middle-Late Triassic. Consequently, we assume that Early Triassic, which is slightly older than the studied monzogranite (245.5±0.7 Ma), might be the best estimate of the formation age of the diabases. The Langshan Early Triassic diabase, together with the recently identified coeval mafic rocks in Wengunshan (Lin et al., 2014), Mandula area (Jian et al., 2010) and Solonker area (Luo et al., 2016; Jian et al., 2010), indicate the presence of the end-Permian to Early Triassic magmatic flare-up in southwestern XMOB.

  • Presence of the Archean and Proterozoic zircon grains (Fig. 5a), together with the positive Zr-Hf-Pb and negative Nb-Ta anomalies in the primitive mantle-normalized diagram of the diabase samples (Fig. 8b), indicate that the parental magma may have undergone crustal contamination during magma ascent. However, all the diabases are mafic with low concentrations of SiO2 (51.13 wt.%-54.55 wt.%) and relatively high Mg# values (43.2-68.2), suggesting that bulk crustal contamination is unlikely. Their low Lu/Yb ratios (0.13-0.14), narrow range of (Th/Yb)N values (5.41-8.39), consistent LREE and LILE distribution patterns (Fig. 8), as well as relatively high εNd(t) values (-3.1 to +1.5), are also inconsistent with significant crustal contamination. Besides, significant crustal contamination would result in linear correlations between Mg# values and εNd(t) values, (Th/Nb)N and Nb/La ratios, but such characteristics are not observed the diabase samples. Thus, crustal contamination, although it cannot be ruled out, does not appear to play a significant role in the petrogenesis of the studied diabases.

    The diabase samples have variable MgO (2.27 wt.%-9.66 wt.%), Ni (24.5 ppm-123 ppm) and Cr (53.6 ppm-494 ppm) contents, indicating that fractional crystallization might have occurred prior to their emplacement. In the Harker diagram (Fig. 7), the positive correlations between MgO and Fe2O3T, Cr and Ni and the negative correlations between MgO and SiO2 indicate considerable degree of fractionation of olivine and/or clinopyroxene. The increasing TiO2 with decreasing MgO, together with the relatively consistent CaO contents imply that the parental magmas only underwent a minor degree of clinopyroxene fractionation (Figs. 7d and 7g). Although Al2O3 is negatively correlated with MgO (Fig. 7b), the negligible Sr and Eu anomalies (Fig. 8) suggest that plagioclase fractionation was insignificant. Moreover, the roughly negative correlations between MgO and TiO2, P2O5 (Fig. 7), along with the negligible Ti anomalies (Fig. 8b), argue against significant fractionation of the Fe-Ti oxides and apatite. Therefore, olivine with a minor amount of clinopyroxene are the major fractionated minerals of the studied diabases.

  • Overall, all the diabase samples have similar trace element and Sr-Nd-Pb isotopic characteristics, so their derivation can be ascribed to a common magma source. The low SiO2 contents (51.13 wt.%-54.55 wt.%) and relatively high MgO, Cr, Ni contents suggest that they were likely originated from mantle- derived components rather than a crustal source (e.g., Rapp et al., 2003). Isotopically, these rocks display relatively high initial ratios of 87Sr/86Sr (0.705 9-0.708 9), with εNd(t) values and Nd model ages respectively ranging from -3.1 to +1.5 (Fig. 10) and 1 473 to 1 034 Ma, implying an enriched lithospheric mantle source. This is also consistent with their relatively high initial ratios of 208Pb/204Pb (35.968-37.346), 207Pb/204Pb (15.448- 15.508) and 206Pb/204Pb (16.280-17.492).

    Arc-like geochemical signatures, such as enrichment in LREE and LILE relative to heavy REE (HREE) and HFSE, are generally related to either extensively crustal contamination or partial melting of an enriched mantle source (e.g., Zhang et al., 2011). As significant crustal contamination can be excluded, the "arc-like" trace element patterns (Fig. 8) and the enriched isotopic signatures (Fig. 10) of the studied diabase were therefore mainly inherited from an enriched lithospheric mantle source. Depletion of HFSE in the primitive mantle normalized diagrams is generally caused by fluid-related and/or melt- related metasomatism in subduction process (e.g., Ma et al., 2014; Duggen et al., 2005; Thirlwall et al., 1994). Thus, we propose that the mantle source of the studied diabase may have been modified by subduction-related fluids and/or melt. Furthermore, the LILE enriched geochemical features are calling for a LILE-enriched mantle source, which means that volatile- bearing minerals (such as phlogopite and amphibole) might be the major host phases in the mantle source origin (e.g., Ionov et al., 1997; Foley et al., 1996). Since melts in equilibrium with amphibole generally have significantly lower Rb/Sr (< 0.1) and higher Ba/Rb (> 20) ratios, whereas melts of phlogopite- bearing source are expected to show significantly lower Ba/Rb ratios (Ma et al., 2014; Furman and Graham, 1999). The relatively low Ba/Rb ratios (2.72-6.56) and variable Rb/Sr ratios (0.08-0.19) of the diabases in this study indicate predominance of phlogopite in their mantle source (Fig. 13a), implying that metasomatism by volatile-rich melts must have occurred prior to the partial melting. Moreover, partial melting in the spinel- stability field generally has low Dy/Yb ratios (< 1.5), while melting in the garnet-stability field shows much higher Dy/Yb ratios (> 2.5) (e.g., Ma et al., 2014). The diabase samples in this study have Dy/Yb ratios ranging from 1.67 to 1.80, indicating that the partial melting was probably operated within the spinel- garnet transition zone lherzolite (Ma et al., 2014; Jiang et al., 2010; Duggen et al., 2005). However, the relatively low Ce/Y (0.97-1.39) and (Tb/Yb)N ratios (1.31-1.45) of the diabase samples indicate that the melts were more likely to be derived from a spinel-stability mantle source (McKenzie and Bickle, 1988). In the Sm/Yb vs. La/Sm and Lu/Hf vs. Sm/Nd diagrams, all the diabase samples plot closed to the 1%-5% partial melting curve of spinel facies (Figs. 13b-13c), indicating that parental magma of the diabases were likely derived from the low degree partial melting of the phlogopite- bearing lherzolite in the spinel-stability field. Compared with the coeval enriched lithospheric mantle derived mafic rocks in the Solonker area (Jian et al., 2010), the diabases in this study have relatively lower εNd(t) values (-3.1-+1.5), older Nd model ages (1 473- 1 034 Ma) and variable initial 87Sr/86Sr ratios (0.705 9-0.708 9) (Fig. 10). Such isotopic characteristics might point to a relatively heterogeneous SCLM (subcontinental lithospheric mantle) source beneath the southwestern XMOB.

    Figure 13.  (a) Ba/Rb vs. Rb/Sr (Furman and Graham, 1999), (b) Sm/Yb vs. La/Sm, and (c) Lu/Hf vs. Sm/Nd diagrams for the studied diabase samples. The melting curves in Figs. 13b and 13c are from Zhao and Asimow (2014).

  • The monzogranite samples in this study are characterized by high Sr/Y (30-45) and La/Yb ratios (Fig. 9), with relatively high Al2O3 (14.15 wt.%-15.98 wt.%) and Sr (310 ppm-380 ppm) contents and low Y (7.16 ppm-10.50 ppm) and Yb (0.60 ppm-1.15 ppm) contents, indicating geochemical characteristics of adakite-like rocks (Martin et al., 2005; Defant and Drummond, 1990). Generally, adakite or adakite-like rocks can be generated by the partial melting of subducted oceanic slab (e.g., Defant and Drummond, 1990), the partial melting of delaminated lower crust (Wang et al., 2006; Kay and Kay, 1993), the partial melting of thickened lower crust (Chung et al., 2003; Atherton and Petford, 1993) and the differentiation of parental basaltic magmas (Castillo, 2012; Prouteau and Scaillet, 2003).

    The studied monzogranites possess high SiO2 (70.09 wt.%-72.63 wt.%), low MgO (0.40 wt.%-0.87 wt.%), Cr (4.74 ppm-8.31 ppm) and Ni (1.15 ppm-6.35 ppm) contents, these features are more akin to the thickened lower crust-derived adakites and metabasaltic and eclogite experimental melts (Fig. 14) (Wang et al., 2006), rather than the subducted oceanic slab- derived or delaminated lower crust-derived adakites (Martin et al., 2005). Their uniform geochemistry, together with the absence of mafic minerals (e.g., amphibole and clinopyroxene), do not support an origin from primary basaltic magma by crustal assimilation and fractional crystallization processes (Castillo et al., 1999). However, the relatively flat HREE patterns (Fig. 8a), as well as the high Zr/Sm ratios (40-64) of these monzogranite samples suggest amphibole-dominated residual phases in the source region, which is consistent with the coeval high Sr/Y granitoids in the Linxi area (Li et al., 2017). Thus, these rocks are unlikely to be originated from a significantly thickened lower crust. Recent partial melting experiments show that adakite-like magmas can also be produced by the low degree partial melting (10%-40%) of a slightly thickened mafic lower crust with P-T condition of 1-1.25 Gpa and 800-950 ℃ (corresponding to a depth of 30-40 km) (Qian and Hermann, 2013). Considering that only moderate thickening of the crust (40-45 km) occurred during the final orogenic evolution in the southern CAOB (Li et al., 2017, 2016a; Zhang et al., 2014b), we propose that the low degree partial melting of the moderately thickened lower crust is likely to be a plausible explanation for the genesis of the studied high Sr/Y monzogranites.

    Figure 14.  (a) MgO vs. SiO2 and (b) Cr vs. SiO2 diagrams for the studied monzogranites. Fields for the subducted oceanic slab-derived adakites, delaminated lower crust-derived adakites, thick lower crust-derived adakites and metabasaltic and eclogite experimental melts (1-4 GPa) are after Wang et al. (2006). The Triassic high-Mg adakitic andesites (Liu et al., 2012) and Early-Middle Triassic high Sr/Y granitoids (Li et al., 2017) in the Linxi area are also plotted for comparison.

    As shown in the εHf(t) vs. U-Pb age diagram (Fig. 12), the Early-Middle high Sr/Y granitoids in northern Linxi have positive zircon εHf(t) values of +7.3 to +15.6, while those in southern Linxi have weak negative to positive zircon εHf(t) values of -1.3 to +5.4 (Li et al., 2017). These characteristics are apparently different from the Middle Triassic high Sr/Y monzogranites in this study, which are characterized by negative zircon εHf(t) values (-25.8 to -8.8) and old TDM2 model ages (2 900- 1 833 Ma). This kind of spatial distribution may indicate that the lower crust source region of the Early-Middle Triassic high Sr/Y rocks in the southwestern XMOB was dominated by ancient crustal materials, the source region south of the Solonker suture zone got a greater contribution from juvenile crustal materials, whereas the area north of the Solonker suture zone was dominated by juvenile components.

  • The final closure of the Paleo-Asian Ocean and the final continental amalgamation/collision were suggested to begin with the initial collision between the SMT and the NCC during the Middle Permian (e.g., Li et al., 2017, 2016a, 2014; Xiao et al., 2015; Jian et al., 2010), leading to the cessation of the arc magmatism and marine sedimentation, accompanied by a short magmatic hiatus and a regional angular unconformity in the XMOB (Li et al., 2016b, 2014; Eizenhöfer et al., 2014). Afterwards, the terminal collision triggered a regional greenschist- blueschist facies metamorphism along the Solonker suture zone in the Late Permian (Li et al., 2017, 2016a; Jian et al., 2010), which marks the onset of the final episode of orogenic evolution in the XMOB.

    As stated above, the Early Triassic diabases in this study were derived from the enriched lithospheric mantle source beneath the southwestern XMOB. Considering that subduction of the Paleo-Asian Ocean had ceased since the Middle-Late Permian, we assume that the slab break-off following the terminal collision was a plausible mechanism for the generation of the studied diabases. Slab break-off can significantly increase the temperature of the overlying lithosphere due to the upwelling of the hot asthenosphere through the slab window, leading to partial melting of the overlying lithospheric mantle, as well as crustal anatexis (e.g., van de Zedde and Wortel, 2001). Likewise, Jian et al. (2010) suggested that the end-Permian to Early Triassic igneous rocks (255-250 Ma) in the Mandula forearc mélange were derived from the decompression melting of mantle peridotite by heat from the upwelling asthenosphere during the slab break-off. Li et al. (2016a) also proposed that the end-Permian to Early Triassic magmatic flare-up in the southeastern Xilinhot area was in response to the final slab break-off in the XMOB.

    The Middle Triassic high Sr/Y monzogranites in this study, combined with the Early-Middle Triassic high Sr/Y granitoids in Linxi area (Li et al., 2017) and central Jilin (Wang et al., 2015a), indicate moderate thickening and shortening of the crust. Thus, considerable contraction must have occurred after the slab break-off, which might be in relation to the ongoing convergence between the NCC and the SMT. The heat from the underplated basaltic magma may lead to: (1) partial melting of the moderately thickened ancient lower crust beneath the southwestern XMOB to generate high Sr/Y magma with negative εHf(t) values, and (2) partial melting of the moderately thickened juvenile lower crust beneath the Solonker suture zone to produce high Sr/Y magma with positive εHf(t) values. In addition, recent studies revealed that the aqueous fluids released from an residual oceanic slab in the upper mantle also had participated in the generation of the Early-Middle Triassic high Sr/Y magmatism along the Solonker suture zone (Li et al., 2017; Song et al., 2015; Zhang et al., 2014b; Liu et al., 2012; Jian et al., 2010).

    The Paleozoic tectonic development of the XMOB was overall dominated by the evolution of the Paleo-Asian Ocean between the NCC and the SMT (e.g., Eizenhöfer et al., 2015, 2014; Xiao et al., 2015, 2003; Windley et al., 2007). During Carboniferous to Early Permian, the progressively double-sided subduction of the Paleo-Asian Ocean enabled an Andean-style continental margins to develop on the northern NCC and the SMT (Liu et al., 2013; Chen et al., 2009; Zhang et al., 2009a; Xiao et al., 2003), accompanied by extensive arc magmatism (e.g., Li et al., 2016b; Chen et al., 2009; Zhang et al., 2009a). The initial collision between the NCC and SMT in Middle Permian prevented further oceanic subduction and terminated the arc-related magmatism. The Paleo-Asian Ocean in the XMOB ultimately closed during the Late Permian, leading to the formation of the Solonker suture zone (Li et al., 2017; Xiao et al., 2015; Jian et al., 2010; Zhang et al., 2009a). The slab break-off of the Paleo-Asian oceanic lithosphere switched the regionally tectonic regime from the prolonged compression to a temporary extension, and then resulted in the magmatic flare-up during end-Permian to Early Triassic (Fig. 15a), such as the Langshan diabase dykes in this study, the E-MORB diabase, sanukitoid, high-Mg diorite and anorthosite in the Solonker ophiolitic mélanges (Jian et al., 2010), as well as those collision-related calc-alkaline granitoids in the southeastern Xilinhot area (Li et al., 2016a). Subsequently, the ongoing weak convergence between the NCC and the SMT initiated moderately crustal thickening and uplift, generating the Early- Middle Triassic high Sr/Y magmas and the high-Mg adakitic andesites (Fig. 15b) (Li et al., 2017; Liu et al., 2012). The post-orogenic extension in the XMOB might be started in the Late Triassic, as evidenced by the 224-208 Ma metamorphic core complex in the Sonid Zuoqi area (Davis et al., 2004), as well as the widespread A-type granitic rocks and alkaline complexes along the northern NCC (Zhang et al., 2014a, 2012). By the end of Early Mesozoic, tectonic evolution of the XMOB was overall controlled by the Mongol-Okhotsk tectonic regime to the northwest and the Paleo-Pacific tectonic regime to the east (e.g., Zhou and Wilde, 2013).

    Figure 15.  Schematic illustrations showing the genesis model and geodynamic environment of end-Permian to Early-Middle Triassic magmatism during the final orogenic evolution of the XMOB. SCLM. Sub-continental lithospheric mantle.

5.   CONCLUSIONS
  • (1) The studied diabases and monzogranites were respectively formed during Early Triassic and Middle Triassic, corresponding to the end-Permian to Early-Middle Triassic magmatic events identified in the southwestern XMOB.

    (2) The Early Triassic diabases were derived from the low degree partial melting of a subduction-metasomatized lithospheric mantle followed by fractionation of olivine and minor clinopyroxene, the Middle Triassic monzogranites were generated by the low degree partial melting of a moderately thickened ancient lower crust.

    (3) The diabases were resulted from the slab break-off of the Paleo-Asian oceanic lithosphere following the terminal collision between the NCC and the SMT in the southwestern XMOB, while the monzogranites were resulted from the subsequently crustal thickening and uplift.

ACKNOWLEDGMENTS
  • This study was jointly supported by the Geological Survey of China (No. 1212011085490) and the National Natural Science Foundation of China (No. 41421002). We sincerely thank the editors and the anonymous reviewers for their critical and constructive comments. We are grateful to Guangqiang Xiong, Hongtao Zhao, Quanliu Chen, and Zhong Wang for their help with the field work. We thank Yu Zhu, Fangyi Zhang and Zezhong Zhang for their laboratory assistance. We also appreciate Yuan Yuan and Yaoyao Zhang for their insightful suggestions. The final publication is available at Springer via https://doi.org/10.1007/s12583-019-1015-5.

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