Introduction

The Altaid Collage (e.g., Sengör et al. 1993; Sengör and Natal’in 1996) or the Central Asian Orogenic Belt (CAOB; e.g., Janh et al. 2000; Xiao et al. 2003; Windley et al. 2007) is one of the largest Paleo-Mesozoic orogenic belts in the world (Fig. 1a). The CAOB lies among the Siberian craton (SC), East European craton (EEC), North China platform (NCP), and Tarim block (TB) and includes continental fragments, low–high T/P metamorphic rocks, ophiolite, continental/oceanic volcanic arc rocks, and accretionary complexes (Petrov et al. 2014). Its geological structure is key to understand the Paleo-Mesozoic geodynamic processes of the Eurasian continent. Various tectonic models for formation of the CAOB were presented. For instance, three major ideas are: (1) the continuous subduction–accretion of oceanic material along a long-lasting arc (Sengör et al. 1993; Sengör and Natal’in 1996); (2) the reworked collision of continental fragments rifted from the supercontinent Rodinia (e.g., Kovalenko et al. 2004); and (3) the time-stepped collision of various island arcs and the subduction of oceanic crust (Badarch et al. 2002; Xiao et al. 2003; Windley et al. 2007) during the amalgamation of the SC, EEC, NPC, and TB.

Fig. 1
figure 1

a Tectonic zoning of North, Central, and East Asia. SC Siberian craton, EEC East European craton, NPC North China platform, TB Tarim block, CAOB Central Asian Orogenic Belt. The rocks of CAOB are divided into following belts and plains: the Sayan-Baikal (SB), Altai-Sayan (AS), North Tien Shan (NTS), Mongol–Okhotsk (MO), Gobi-Khingan (GK), South Tien Shan (STS), Junggar-Balkhash (JB), Central Kazakhstan (CK), Ob-Zaysan (OZ), East Ural (EU), Hangai-Daur (HD), Solonker (S) Beishan (B), and Sikhote-Alin (SA) belts and the Amur-Zeya (AZ), Songliao (Sl), Junggar (J), Turan (T), and West Siberian (WS) plains (simplified after Petrov et al. 2014). The Main Mongolian Lineament (MML) marks the border between the AS, SB, and HD belts and the GK belt and is considered as one of the major tectonic lines of the CAOB. b Geological outline of North Mongolia–West Transbaikal region. This region is represented by the black square in the middle a. The terranes in this region are as follows: Central Mongol (CM), Tuva Mongol (TM). Dzhida (Dz), Khamar Daban (KhD), Tunka (Tu), Ikat (Ik), and Eravna (Er) of the SB belt and Zag (Zg), Tsetserleg (Tt), Harhorin (Hh), Haraa (Hr), Asralt Hairhan (AH), Daur (Dr), Aga (Ag), Onon (On), and Ulaanbaatar (Ub) of the HD belt (after Bulgatov and Gordienko 1999; Petrov et al. 2014; Tomurtogoo 2012). The number reflects the locality (L) of the Late Paleozoic folds and thrusts: L.1 the Züünharaa area; L.2 the Zuunmod area; L.3 the Bayangol area; L.4 the Zaamar area; L.5 the Tunka area; L.6 the Ul`zutui, Oldynda and Kydzhimit areas; L.7 the Bagdarin area; and L.8 the Ulaanbaatar area (compiled by this study with Purevsuren and Narantsetseg 1991; Dejidmaa 2003; Zhimulev et al. 2011; Gordienko et al. 2012; Ruzhentsev et al. 2012; Takeuchi et al. 2012; Altanzul and Baasandolgor 2014). The study area (black dashed box) is located within the Haraa terrane (Hr), from Fig. 2, in greater detail

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Fig. 2
figure 2

a Geological map around the Züünharaa area. Red lines are low-angle southward-dipping thrust: (1) conformed; (2) concealed; and (3) inferred. Circle points are the location of samples with zircon U-Pb age. Triangle points are the location of outcrops, which are shown in Figs. 3a–d, 6a–d, and 7a–d. b Geological cross section along AA’. c Bedding planes of rocks in the Haraa Group, Ulaan Öndör Formation, and Örmögtei Formation by pole on the Stereonet diagram

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The North Mongolia–West Transbaikal region (Fig. 1a, b), which lies north of the Main Mongolian Lineament (MML; Badarch et al. 2002; Tomurtogoo 2006; Kröner et al. 2007; Windley et al. 2007), is one of the significant portions for Paleozoic–Mesozoic tectonics of the CAOB in the Northeast Asia. Geologically, this region is comprised of rocks from the Middle Devonian Sayan-Baikal and Middle–Late Triassic Hangai-Daur belts (Fig. 1b, Petrov et al. 2014). The Paleozoic accretionary complexes (Kurihara et al. 2009; Tomurtogoo 2012) in the Hangai-Daur belt were formed by the northward subduction of the oceanic plate (Mongol Okhotsk oceanic plate; e.g., Tang et al. 2016) beneath the SC (Zorin 1999; Donskaya et al. 2013). The Mongol–Okhotsk Ocean was previously present between the SC and NCP/TB (Fig. 1a, e.g., Enkin et al. 1992; Gordienko 2001; Badarch et al. 2002; Donskaya et al. 2013; Ruppen et al. 2013). However, the closing process and termination of the Mongol–Okhotsk Ocean are still unclear (e.g., Zorin 1999; Kravchinsky et al. 2002; Kelty et al. 2008; Donskaya et al. 2013). Only the closure of the ocean, regarded to be the result of the collision between the SC and the NCP, is generally agreed upon (e.g., Enkin et al. 1992; Kravchinsky et al. 2002; van der Voo et al. 2015; Fritzell et al. 2016).

Studying the geological processes and features from the Paleozoic accretionary complexes to the continental margin of the SC in the North Mongolia–West Transbaikal region will facilitate understanding of the tectonics of the CAOB (Fig. 1b). However, the structural and tectonic aspects in this region remain poorly understood because of a lack of detailed tectonic studies. 3D imaging and structural interpretation are an essential part of establishing the regional tectonic setting of an area. This paper describes the lithology, stratigraphy, geological structure, and U-Pb age of the rocks in the Züünharaa area, about 150 km north of Ulaanbaatar city, Mongolia (Fig. 1b). The Züünharaa area is located within the Haraa terrane of the Hangai-Daur belt, which has a continental affinity. The tectonic implication of the low-angle south-dipping thrust in this area is discussed within the North Mongolia–West Transbaikal region.

Geology

The CAOB (Fig. 1a) is composed of Middle Devonian (Sayan-Baikal, SB; Altai-Sayan, AS; Kazakhstan, K; North Tien Shan, NTS), Late Carboniferous–Permian (Mongol–Okhotsk, MO; Gobi-Khingan, GK; South Tien Shan, STS; Junggar-Balkhash, JB; Central Kazakhstan, CK; Ob-Zaysan, OZ; East Ural, EU), Middle–Late Triassic (Hangai-Daur, HD; Solonker, S; Beishan, B) and Early Cretaceous (Sikhote-Alin, SA) belts and Mesozoic–Cenozoic (Amur-Zeya, AZ; Songliao, Sl; Junggar, J; Turan, T; West Siberia, WS) plains (Petrov et al. 2014). The basement rocks of Mongolia are divided into two regions, North and South, separated by the Main Mongolian Lineament (Fig. 1a; MML; e.g., Badarch et al. 2002; Tomurtogoo 2006; Kröner et al. 2007; Windley et al. 2007). The North Mongolia–West Transbaikal region (Fig. 1b), which lies north of the MML, is divided into the Sayan-Baikal (SB) belt and Hangai-Daur (HD) belt (Fig. 1a, b; Petrov et al. 2014). The SB belt is considered to be a collage of amalgamated blocks that have been accreted to the Siberian craton, while the HD belt is regarded as accretionary complexes with shelf facies rocks (e.g., Kurihara et al. 2009; Tomurtogoo 2012, 2014; Bulgatov and Gordienko 2014).

The SB belt is mainly composed of Precambrian continental basement rocks, Neoproterozoic–Carboniferous sedimentary rocks and Permian–Triassic volcanic–plutonic rock complexes (Bulgatov and Gordienko 1999, 2014; Parfenov et al. 2009; Donskaya et al. 2013). The rocks of SB belt are divided into following seven geological units (Fig. 1b): Central Mongol, Tuva Mongol, Dzhida, Khamar Daban, Tunka, Ikat and Eravna terranes (Parfenov et al. 2004a, 2009; Tomurtogoo 2006, 2012, 2014). Paleoproterozoic to Early Cambrian metamorphic rocks, ophiolitic rocks, and volcanic-sedimentary rocks are largely exposed in the Central Mongol, Tuva Mongol, Dzhida, and Khamar Daban terranes. The Paleozoic rocks occur in the Tunka, Ikat, and Eravna terranes (Parfenov et al. 2004a, 2009). The Permian–Triassic volcanic–plutonic rock complexes overlap the rocks of the SB belt and form the Selenga superimposed volcano–plutonic belt (Fig. 1b; Bulgatov and Gordienko 1999, 2014; Parfenov et al. 2004b, 2009).

The Central Mongol and Tuva Mongol are considered to be Precambrian continental fragments, while the Dzhida and Khamar Daban terranes are regarded to be Early Paleozoic accretion–collision orogens with a complex N-verging fold-thrust structure (e.g., Tomurtogoo 2006; Belichenko et al. 2003; Gordienko et al. 2007). The Tunka terrane is located on the west side of the southern rim of the Siberian craton (Fig. 1b). This terrane consists mainly of Ordovician–Silurian terrigenous-volcanogenic-carbonate sedimentary rocks and is intruded by Tunka granite of the Sarkhoi plutonic complex with an U-Pb age of 462.6 ± 7.8 Ma (Zhimulev et al. 2011). The molasses-like Late Devonian–Early Carboniferous Sagan-Sair Formation overlaps the Tunka granite and Ordovician–Silurian sedimentary rocks (Buslov et al. 2009; Zhimulev et al. 2011). The rocks of the Tunka terrane form a complex N-verging fold-thrust structure (e.g., Ryabinin et al. 2011; Zhimulev et al. 2011).

The rocks of the Ikat and Eravna terranes occur as scattered fragments or variable size of xenoliths in the Angara-Vitim batholith (Donskaya et al. 2013) of the Permian–Triassic superimposed volcanic–plutonic rock complexes (Fig. 1b, e.g., Parfenov et al. 2009; Badarch 2005; Ruzhentsev et al. 2012). The Ikat and Eravna terranes are primarily composed of Neoproterozoic–Carboniferous weakly metamorphosed carbonate-terrigenous deposits with minor volcanogenic rocks (Mazukabsov et al. 2010). The Neoproterozoic–Carboniferous rocks form a complex N-verging fold-thrust structure in the Ul’zutui, Oldynda, Kydzhimit areas of the Eravna terrane and the Bagdarin area of the Ikat terrane (Ruzhentsev et al. 2006, 2012).

The HD belt is largely composed of Cambrian–Ordovician proximal shallow marine sedimentary rocks and Paleozoic accretionary complexes (Tomurtogoo 2006, 2012; Kurihara et al. 2009). The Paleozoic rocks of the HD belt are divided into the following nine geological units (Fig. 1b): Zag, Tsetserleg, Harhorin, Haraa, Asralt Hairhan, Daur, Aga, Onon, and Ulaanbaatar terranes (Bulgatov and Gordienko 1999, 2014; Tomurtogoo 2012, 2014). Most terranes in the HD belt are regarded as accretionary complexes, with the exception of the Zag and Haraa terranes. The Ulaanbaatar terrane (late Devonian Early Carboniferous accretionary complex) is composed mainly of basalts, limestone, Silurian–Devonian radiolarian chert, siliceous mudstone, and clastic rocks (e.g., Kurihara et al. 2009; Nakane et al. 2012; Suzuki et al. 2012; Takeuchi et al. 2012; Tsukada et al. 2013). The Ulaanbaatar, Harhorin, and Tsetserleg terranes are structurally and lithostratigraphically correlative (Kurihara et al. 2009; Purevjav and Roser 2013; Tsukada et al. 2013). The Asralt Hairhan terrane is considered to have a metamorphic affinity to the Ulaanbaatar terrane (e.g., Tomurtogoo 2012; Gordienko et al. 2012). The Daur, Aga, and Onon terranes consist of accretionary complexes dominated by deformed–metamorphosed mélanges containing fragments of ophiolite, arc-related volcanic, and sedimentary rocks with minor amounts of radiolarian chert and limestone (Bulgatov and Gordienko 1999; Zorin 1999; Badarch 2005; Tomurtogoo 2012).

The Zag terrane consists of highly deformed pelitic and psammitic schist of the Zag Group, which yielded K-Ar ages of 459.9 ± 9.1 and 447.4 ± 9.0 Ma (Kurimoto et al. 1998; Badarch 2005). The Haraa terrane (Fig. 1b) is largely composed of Cambrian–Lower Ordovician metamorphosed clastic rocks of the Haraa Group (Tomurtogoo 2006, 2012). The rocks of the Haraa Group are intruded by the Boroogol plutonic rock complex giving the U-Pb age of 460–440 Ma (e.g., Kröner et al. 2007; Hou et al. 2010). The Haraa Group and Boroogol plutonic rock complexes are unconformably overlain by Devonian–Permian volcanic and clastic rocks (Badarch et al. 2002). The younger Paleo-Mesozoic plutonic rocks intrude into the above sedimentary and igneous rocks (Tomurtogoo 2006; Donskay et al. 2013).

The bedding and structural planes in the Zag and Tsetserleg terranes generally trend northwest and steeply dip north- or southward, west of the HD belt (e.g., Badarch et al. 2002; Buchan et al. 2001; Purevjav and Roser 2013; Tsukada et al. 2013), whereas those in the Haraa, Asralt Hairhan, Ulaanbaatar, Daur, and Aga terranes trend northeast and steeply dip northward in the east (e.g., Zorin 1999; Kurihara et al. 2009; Takeuchi et al. 2012). The rocks of the Harhorin terrane trend north-northwest and dip southwestwardly in fault contact with the Tsetserleg terrane in the west by the northwestwardly trending fault system (Fig. 1b; Tseden et al. 1992; Tsukada et al. 2010; Purevjav and Roser 2013). The rocks of the Onon and Aga terranes form an S-like bowing structure, elongating to northeast and dipping northward (Fig. 1b, e.g., Zorin 1999). The rocks of the Asralt Hairhan and Ulaanbaatar terranes form southeast-verging composite folds associated with northward-dipping thrusts (e.g., Kurihara et al. 2009; Gordienko et al. 2012; Nakane et al. 2012; Suzuki et al. 2012; Takeuchi et al. 2012). The Asralt Hairhan terrane is in fault contact with the Haraa terrane via the northeastwardly trending Late Mesozoic Yeroogol sinistral strike-slip fault system at its northern end (Fig. 1b; Tseden et al. 1992; Kotlyar et al. 1998; Altanzul and Baasandolgor 2014).

The study area (L.1, Fig. 1b) is located within the Haraa terrane and includes Züünharaa Town, Mongolia. The study area exposes metamorphosed clastic rocks of the Haraa Group, volcanic rocks of the Ulaan Öndör Formation, and clastic rocks of the Örmögtei Formation in ascending order (Figs. 2, 3). The rocks of the Haraa Group are intruded by granitic rocks. A part of the granitic rock is monzogranite according to the I.U.G.S. classification of plutonic rocks (Streckeisen and Le Bas 1991). The granitic rock gradually changes into finer rhyolitic rocks in places and is intruded by later dikes of diorite and dolerite (Figs. 2, 3). Tomur et al. (1994) defined the volcaniclastic rocks in the Zuunmod area, 30 km south from the study area (L.2, Fig. 1b), as Devonian Ulaan Öndör Formation. The volcaniclastic rocks in the study area are regarded as equivalent to the Ulaan Öndör Formation based on their lithological similarity (Purevsuren and Narantsetseg 1998). The clastic rocks of the Örmögtei Formation unconformably overlie the granitic rock with basal conglomerate (Figs. 2a, b, 3d).

Fig. 3
figure 3

Generalized lithostratigraphic column with field occurrence of distributed rocks in the Züünharaa area. a Rocks are as follows: (1) metamorphosed sandstone–mudstone; (2) equigranular granite; (3) porphyritic granite; (4) porphyritic rhyolite; (5) diorite–dolerite dikes; (6) dacitic to rhyolitic lava; (7) tuff breccia; (8) clast-supported conglomerate; (9) arkosic sandstone; (10) matrix-supported conglomerate; (11) bedded sandstone; and (12) alternating beds of mudstone and sandstone. b Dark-light banded beds of the metamorphosed sandstone–mudstone showing a stratigraphic up (black arrow). c Xenolith from the Haraa Group in porphyritic granite shows evidence of intrusion. d Angular to sub-rounded clasts of rhyolitic to dacitic lava, porphyritic granite–rhyolite (Rh), and the metamorphosed sandstone–mudstone (Ms) are included in the tuff breccia, indicating conformity. e Equigranular granite is intruded by dolerite dike and is unconformably overlain by basal clast-supported conglomerate; and the co-author pinpoints the unconformity. The length of a mechanical pencil, a hammer, and the height of the co-author denote 14.5, 20, and 170 cm, respectively

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Methods and samples

We carried out field research on a 600 km2 area in and around Züünharaa Town to detail the lithology, stratigraphy, and geological structure (Figs. 2, 3, 4, 5, 6, 7, 8). The obtained data are used to examine tectonics of the SB and HD belts. Based on the obtained data, the geological map and cross section are newly established.

Fig. 4
figure 4

Photomicrographs of thin sections from distributed rocks in the study area. a Metamorphosed sandstone from the Haraa Group. Abundant hornblende, biotite, and episode occur in a finer matrix of fine-grained sandstone. b Equigranular granite. Magmatic zircons are included in the biotite, aplite, and hornblende. c, d Bedded sandstone from the Örmögtei Formation. Abundant lithic fragments are derived from igneous rocks and are included in the bedded sandstone. Abbreviations on photomicrographs: quartz (Qz), K-feldspar (Kfs), plagioclase (Pl), hornblende (Hbl), biotite (Bt), epidote (Ep), apatite (Ap), zircon (Zr), lithic fragment (Lf), plane-polarized light (ppl), AND cross-polarized light (xpl). The white line denotes 1 mm

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Fig. 5
figure 5

Petrographic classification and provenance of sandstone from the Örmögtei Formation. a Q–F–L diagram from Okada 1971 and b Qt–F–L from Dickinson et al. (1983). Sandstone from the Örmögtei Formation is classified as feldspathic arenite and are plotted in the field of “basement uplift” and “volcanic arc”. Carboniferous sandstone results of Bussien et al. (2011), Hara et al. (2013), Ruppen et al. (2013), and Suzuki et al. (2012) are plotted for comparison. Total quartz (mono- and polycrystalline grains; Qt), feldspars (plagioclase and potassium feldspar; F), and lithic fragments (L)

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Fig. 6
figure 6

Occurrence of cataclasite rocks in and around of the low-angle southward-dipping thrust at different scales. Boxes are the approximate location of the following figures. Scale lines denote ca. 100 m, 5 cm, 1 cm, and 0.5 mm, for views a–d, respectively. a Porphyritic rhyolite overthrust onto the Haraa Group to form a cataclasite zone, more than 40 m wide at Shirhentseg hill, east of Züünharaa Town. b Rock outcrop shows an intensively sheared part of the cataclasite zone; the dashed line marks a lithological boundary between the porphyritic rhyolite and the metamorphosed sandstone of the Haraa Group. c Polished surface of a rock sample of non-foliated cataclasite along an oriented cut. d Photomicrographs show an angular clast, which is derived from porphyritic rhyolite. Abbreviations on photomicrographs: quartz (Qz), K-feldspar (Kfs), plagioclase (Pl), plane-polarized light (ppl), and cross-polarized light (xpl)

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Fig. 7
figure 7

Representative outcrops of foliated cataclasite from the low-angle southward-dipping thrust. Sub-horizontal cataclasites exposed in many places. Locations are shown in Fig. 2a and in Table 1. Solid lines represent Y-planes, and dashed lines are P-planes. Shear sense (arrow) was determined from the Y–P composite planar planes. The great circles in the lower hemisphere projection show the Y–P planes of the outcrop. a Rock outcrop of metamorphosed sandstone from the Haraa Group shows a sub-horizontal fault plane. b Sub-horizontal cataclasited section of the tuff breccia shows few Y-planes. Outcrops c, d are foliated cataclasite derived from porphyritic granite–rhyolite showing several Y–P planes, which suggest top-to-the-north shearing

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Fig. 8
figure 8

Orientation data for all Y–P planes measured from several outcrop surface exposures of the low-angle southward-dipping thrust in the study area. a Rose diagrams by (1) data from the strike; (2) dip direction; and (3) dip angle of the Y–P planes. b Stereonet diagram by (4) data from poles of the Y–P planes. The Y-planes and P-planes are represented by blue pluses and red dots, respectively. c Stereonet diagram by (5) data from the main thrust plane. The great blue circle (solid) is the main Y-plane, and the great red circle (dashed) is the average P-plane. The main Y-plane and average P-plane (5) suggest top-to-the-north shearing. The orientation data refer to Table 1

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Table 1 Orientation data of the Y–P planes from the low-angle southward-dipping thrust in the Züünharaa area
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Zircon LA–ICP–MS dating of the granitic rock was carried out for a sample from equigranular granite (#12Hr1601) and two samples from porphyritic granite–rhyolite (#11Hr54 and #12Hr1011) to constrain geological framework of the study area (Fig. 2a). Sample preparation for U-Pb dating of zircons by LA–ICP–MS is described in Chang et al. (2006). The zircon samples were concentrated using the conventional mineral-separation techniques, including crushing and pulverizing, followed by separation (by hand) using magnets. These zircon grains were then mounted in an epoxy resin, and diamond polished to expose the interior. To investigate the internal structure of the individual grains, a scanning electron microscope (SEM; Hitachi S-3400N equipped with Gatan MiniCL) installed at Nagoya University was used to obtain backscattered electron (BSE) and cathode luminescence (CL) images.

The analyzed zircon grains are colorless, 50–100 µm in length along their major axes and have a length-to-width ratio of 1.5–2. These zircon grains mostly occur as subhedral to euhedral prisms, and contain distinct cores with concentric zoning, consistent with crystallization from a viscous magma, and overgrowths (Fig. 9a–d, e.g., Corfu et al. 2003).

Fig. 9
figure 9

Representative cathodoluminescence (CL) images of zircon grains from dated granitic rocks. Zircons from sample a #12Hr1601 (non-deformed equigranular granite), b #11Hr54 (fractured porphyritic granite), and c #12Hr1011 (a clast in the cataclasite derived from the porphyritic rhyolite). The locations of the analytical spots (C core and R rim) are shown as circles together with grain numbers (G#) and Concordia ages (in Ma). Diameter of analytical slot is 25 micrometer, and white scale denotes 100 µm. The analytical results refer to Table 2

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Table 2 Analytical results of LA–ICP–MS U-Pb dating for granitic rock from the Züünharaa area
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The samples were analyzed by Inductively Coupled Plasma Mass Spectrometry (Quadrupole type ICP-MS; Agilent 7700x), which was connected with the NWR-213 LA system (Electro Scientific Industries, Inc.) installed at Nagoya University. The ablation pit size was 25 µm under conditions of 10 Hz repetition rates with energy densities of ~12 J/cm2. Materials used for calibration were 91,500 standard zircon (1062.4 Ma; Wiedenbeck et al. 1995) and silicate glass reference materials produced by the National Institute of Standards and Technology (NIST): SRM 610 (Horn and Von Blanckenburg 2007). Further details on the accuracy and reproducibility of U-Pb can be found in Orihashi et al. (2008), Iwano et al. (2013) and Kouchi et al. (2015).

Detailed geology of the study area

Lithology and stratigraphy

(1) Haraa Group

The rocks of the Haraa Group consist of metamorphosed sandstone with mudstone intercalations (Fig. 3a, b). Sandstone is usually fine- to medium-grained, sub-angular, generally well-sorted but poorly sorted in some horizons. It includes many grains of quartz and feldspar in a finer matrix. Hornblende, amphibole, biotite, chlorite, and epidote occur in the rocks as metamorphic minerals (Fig. 4a). The alternating beds of sandstone and mudstone show sedimentary structures, such as graded bedding (Fig. 3b), cross bedding, and load casts in some places. The thickness of a bed varies from several millimeters to several meters. Several xenoliths, which are derived from metamorphosed sandstone–mudstone, are observed in the granitic rocks (Fig. 3c). The total thickness of the metamorphosed clastic rocks exceeds 2000 m (Fig. 3a).

(2–4) Granitic rocks and (5) mafic dikes

In this study, the granitic rock is subdivided into three types based on texture: (2) equigranular granite (monzogranite), (3) porphyritic granite, and (4) porphyritic rhyolite (Figs. 2a, 3a). The equigranular granite generally includes K-feldspar and plagioclase as major minerals, and minor amounts of quartz, hornblende, biotite, apatite, and zircon (Fig. 4b). The major minerals are 1–3 mm in size. Subhedral to euhedral K-feldspar shows micro-perthite texture and twinning with sericite inclusion. Plagioclase shows that clear stripe twining is often turbid in the interior, owing to alteration (Fig. 4b). Hornblende and biotite, displaying brown under plane-polarized light, occur as slender needles or lathes.

The porphyritic granite is inter-gradual with porphyritic rhyolite, and the boundary between them is obscure (Fig. 2a, b). Porphyries of quartz, K-feldspar, and plagioclase, up to 7 mm in size, are embedded in a finer groundmass in the porphyritic granite. Minor amounts of amphibole, biotite, and sericite are also included in the porphyritic granite. Some porphyritic granite has abundant crystals of sphene, zircon, and opaque minerals. The porphyritic rhyolite includes phenocrysts of quartz and feldspar in a microcrystalline groundmass. Phenocrysts are up to 3 mm in size. Biotite and hornblende, up to 0.5 mm in size, also occur in a groundmass.

Small dikes of (5) diorite and dolerite intrude into the porphyritic granite and equigranular granite (Figs. 2a, 3e). The diorite is composed of plagioclase, hornblende with a small amount of K-feldspar, biotite, and sphene. The dolerite includes phenocrysts of plagioclase, augite and quartz in a finer groundmass of intersertal plagioclase.

(6–7) Ulaan Öndör Formation

The rocks of the Ulaan Öndör Formation overlie the rocks of the Haraa Group and the granitic rocks (Figs. 2a, b, 3a). This formation is composed mainly of (6) rhyolitic–dacitic varicolored lava and (7) tuff breccia, in ascending order (Fig. 3). The lava includes irregular-shaped phenocrysts of quartz and feldspar in a cryptocrystalline groundmass. The tuff breccia is yellowish green, clast-supported and poorly sorted, and intercalates fine- to medium-grained tuffaceous mudstone layers in some horizons. Tuff breccia includes angular to sub-rounded clasts of rhyolite–dacite, porphyritic granite–rhyolite, and metamorphosed sandstone–mudstone (Fig. 3c). The clasts are up to 10 cm in size. Tuffaceous mudstone rarely contains quartz and plagioclase in a finer matrix. The total thickness exceeds 300 m (Fig. 3a).

(8–12) Örmögtei Formation

The clastic rocks of the Örmögtei Formation unconformably overlie the equigranular granite (Figs. 2a, b, 3a, e). This formation is subdivided into the lower and upper members. The lower member consists of (8) clast-supported conglomerate and (9) medium- to coarse-grained massive arkosic sandstone and the upper member consists of (10) matrix-supported conglomerate, (11) fine- to medium-grained bedded sandstone, and (12) alternating beds of mudstone and sandstone, in ascending order (Fig. 3a).

Clast-supported conglomerate includes abundant boulders, up to 5 m in size, of porphyritic granite–rhyolite. Arkosic sandstone is largely composed of quartz and feldspar grains and includes clasts of cobble- to pebble-sized, rounded porphyritic rhyolite.

The matrix-supported conglomerate overlies the arkosic sandstone. The matrix-supported conglomerate intercalates the upper part of the bedded sandstone. The bedded sandstone dominated by quartz, plagioclase, potassium feldspar, and rock fragments (Fig. 4c, d) is overlain by the alternating beds of mudstone and sandstone. The sandstone of the alternating beds is composed largely of quartz, plagioclase and potassium feldspar, and rock fragments. Each bed is several centimeters thick. Sedimentary structures, such as graded bedding and cross lamina in the alternating beds, show stratigraphic up (Fig. 2a). The sandstone is classified as feldspathic arenite in the Q–F–L diagram (Fig. 5a; Okada 1971). In the Qt–F–L diagram (Dickinson et al. 1983), this sandstone is plotted in the field of “basement uplift” and “volcanic arc” (Fig. 5b). The sandstone has clasts of rhyolitic tuff and also includes lesser amounts of zircon, muscovite, biotite, epidote, and chlorite (Fig. 4c, d). Some mudstone layers in the alternating beds yield brachiopod fossils. The type section of the Örmögtei Formation, 45 km north from the study area, is assigned to the Visean based on a flore and a brachiopod assemblage, Tomiodendron ex gr.kemerovience and Dyscritella mergensis–Lanopora mongolica respectively (e.g., Ariunchimeg 2011; Tolokonnikova et al. 2014). Therefore, it is likely that the Örmögtei Formation in this area is also Visean. The total thickness of this formation exceeds 1450 m (Fig. 3a).

In conclusion, the study area can be divided into four types of rocks, i.e., the metamorphosed clastic rocks of Haraa Group, granitic rocks with mafic dikes, the volcanoclastic rocks of the Ulaan Öndör Formation, and the clastic rocks of the Örmögtei Formation (Figs. 2a, 3a), which are subdivided into 12 distinguishable lithological units (Fig. 3a). The Haraa Group intruded via granitic rocks and unconformably covered the Ulaan Öndör and Örmögtei formations. The relationship between the Ulaan Öndör Formation and the Örmögtei Formation is not clear due to poor exposure in the study area, and is inferred to have conformity based on Tomur et al. (1994).

Structure

The rocks of the Haraa Group are exposed in the western and eastern parts of the study area (Fig. 2a). The bedding planes of the rocks trend northwesterly and steeply dip southward to take a monoclinal structure in the western part (Fig. 2c). In the eastern part, the southern side strikes N40°–60°E and dips 40°–65°N, and the northern side strikes N40°W–N10°E and dips 45°–80°N to take a complex fold structure (Fig. 2a, c). The rocks of the Ulaan Öndör formation occur in several small areas at central and western parts of the study area (Fig. 2a). The bedding planes of the volcanoclastic rocks strike N60°E and dip 35°–45°N (Fig. 2a, c). The Örmögtei Formation is open folded with a low-angle axis plunging westward (Fig. 2a). The southern wing strikes N 30°–50°E and dips 15°–45°N, and the northern wing strikes N 45°–60°W and dips 15°–35°S to take a syncline structure (Fig. 2a, c).

We identified a regional low-angle thrust trending E–W and dipping gently southward. The thrust cuts all geologic units throughout in this area with the exception of the Örmögtei Formation (Fig. 2a, b). The rocks in and around the thrust were intensely fractured, forming a cataclasite zone (Fig. 6a). A cataclasite zone with a maximum width of 40 m extends for at least 30 km along the W–E and 20 km along the S–N and exposes from the 900 to 1000 m elevations (Figs. 2a, 6a).

Abundant angular clasts derived from the host rocks, such as meta-sandstone, porphyritic rhyolite–granite, volcaniclastic rocks, and others, are included in a finer matrix, demonstrating the random fabric of the cataclasite (Fig. 6b, c). The cataclasite is rarely foliated in places but generally non-foliated. The clasts are formless, lenticular, spherical, or tabular, and vary in size from several millimeters to several centimeters (Fig. 6c, d). The clasts (>2 mm) make up to 60% of total the volume (Fig. 6c). The cracks in the clasts are filled by grains fed from the matrix. The matrix of the cataclasite is entirely composed of cryptocrystalline minerals (Fig. 6d) derived from host rocks and minor amounts of secondary chlorite, calcite, and sericite (excluding epidote) which suggests a formation temperature above 200 °C (Henley and Ellis 1983). The matrix, which includes cement, occupies more than 40% of the total volume. According to the classification of fault rock by Woodcock and Mort (2008), the rocks in and around the thrust in this area are chaotic breccia and protocataclasite. Textures formed by cataclastic flow and pressure solution are usually observed, and no evidence of ductile deformation, such as grain boundary migration, sub-grain rotation, bulging, and undulose extinction, is observed. Taking into consideration that cataclasites generally occur at conditions lower than 4 kbar (Passchier and Trouw 1998), the present cataclasite without epidote seems to have been formed under conditions of less than 200 °C and 4 kbar.

The foliated cataclasite shows composite planar structures, such as Y- and P-planes of subsidiary fractures or so-called Riedel shears (Riedel 1929). The Y- and P-planes (e.g., Tchalenko and Ambraseys 1970; Bartlett et al. 1981) are defined by thin layers, several millimeters wide, of the clay minerals in parallel orientations (Fig. 7a–d; similar to observations from other brittle shear zones in Mukherjee 2012, 2013a, 2014, 2015). The respective spacing of the Y- and P-planes is ~40 and ~10 cm. The Y-and P-planes generally strike N40°W–N10°E and N50°W–N30°E and dips 10°–35° and 30°–60° (Table 1; Fig. 8a). The assembled data of the Y-plane show generally sub-horizontal planes (Fig. 8b). It suggests that the main fault plane strike ~N18°W and dip 9°S in the study area (Fig. 8c). The average P-plane strikes N76°E and dip 27°S (Fig. 8c). The angle between Y- and P-planes is ~20° (Figs. 7c, d, 8a). Shear sense is determined from the Y–P planes (as in Mukherjee 2010a, b; Mukherjee and Koyi 2010a, b; Misra and Mukherjee 2016; Babar et al. 2016; Kaplay et al. 2016). Shear heating (e.g., Mukherjee 2017) was not manifested in these shear zones (Mukherjee and Mulchrone 2013; Mulchrone and Mukherjee 2015, 2016). The Y–P planes in the foliated cataclasite indicate top-to-the-north shear (Figs. 7c, d, 8c). The cataclasites in the northwestern area discontinue in the east, and the clastic rocks of the Örmögtei Formation were not affected by shearing (Fig. 2a). Two younger NW-trending faults are recognized in the central and eastern part of the study area. These high-angle or vertical faults cut both the non-foliated and foliated cataclasite near Mt. Sevsüül and Mt. Zürüü (Fig. 2a).

Summarizing the above observations of rock structure: the bedding planes of the Haraa Group chiefly form a complex fold structure (Fig. 2c); the rocks of Ulaan Öndör Formation have NE-trending, NW dipping bedding planes (Fig. 2c); and the rocks of Örmögtei Formation form an open syncline fold (Fig. 2a, c). We first discovered a low-angle, southward-dipping thrust in this area, which has a top-to-the-north shearing (Fig. 8b, c). The thrust clearly cuts the Haraa Group, the granitic rocks, and the Ulaan Öndör Formation, with the exception of the Örmögtei Formation (Fig. 2a, b).

Age of the granitic rocks

As a result of the BSE and CL examination, a total of 124 zircon grains, with an exposure of more than 40 µm in diameter on the epoxy resin, were chosen for analysis from samples #12Hr1601 (equigranular granite), #11Hr54 (porphyritic granite), and 12Hr1011 (porphyritic rhyolite). Of these, 16 grains were ablated twice at the core and rim. Data were obtained from a total of 140 spots (Table 2) and plotted in a Concordia diagram using the Isoplot 3.75 software (Ludwig 2012). A total of 77 age data points were excluded from examination for various reasons (e.g., less than 0.05 Concordia age probability, or ages with an error of >100 Ma). The remaining 63 data points were used in this study and plotted on a weighted mean diagram (Fig. 10a–c). It is generally accepted that zircon grains with a high Th/U ratio (>0.2) are of magmatic origin, whereas those with a low Th/U ratio (<0.1) have undergone a secondary process, such as metamorphism and hydrothermal alteration (Hartmann et al. 2000; Hartmann and Santos 2004). The zircon samples in the present study, showing a high Th/U ratio (0.21–0.79), can thus be considered magmatic in origin (Table 2).

Fig. 10
figure 10

U-Pb Concordia and mean weighted diagrams of zircon grains from granitic rocks. a Equigranular granite, b porphyritic granite, and c porphyritic rhyolite. The analytical results refer to Table 2. n A total number of analytical spot. The error ellipses in the Concordia plots are 2σ

Full size image

Sample #12Hr1601

This sample was collected from non-deformed equigranular granite (Fig. 2a). It is coarse-grained and holocrystalline dominated by microcline, plagioclase, hornblende, and biotite with minor zircon. The zircon grains of the equigranular granite are mostly light brown, subhedral to euhedral, semi-transparent, and up to 200 µm in length along their major axis. Most zircons have a few inclusions and cracks and yield dark unzoned CL images (e.g., Fig. 9a; Grain number 46 or G#46).

The Th/U ratio is 0.21–0.5 (Table 2). The ages from 31 concordant analyses ranging from 561 to 480 Ma (Fig. 10a) were grouped at 540.9 ± 8 Ma (10 grains, MSWD = 4.9, probability = 0.026) and 498.3 ± 3.3 Ma (21 grains, MSWD = 0.43, probability = 0.51). The older ages (~540 Ma) are likely to have been obtained from inherited cores within the zircon samples. The younger ages (~498 Ma) are likely to indicate the age of the magmatic event of the equigranular granite (Fig. 10a; Table 2).

Sample #11Hr54

This sample was taken from fractured porphyritic granite near the thrust (Fig. 2a). It is coarse-grained and holo- to hypocrystalline dominated by quartz, K-feldspar, and plagioclase with minor sphene, zircon, and opaque minerals. The zircons from this sample are yellow to colorless, transparent, euhedral to subhedral, up to 150 µm in length along their major axis, and have a length-to-width ratio of 1:2 (Fig. 9b). The zircons have irregular cracks and commonly include apatite needles (e.g., Fig. 9b, G#36 and G#37). Some zircon grains show inherited cores with concentric oscillatory zoning in their CL images (Fig. 9b, G#16).

The Th/U ratio is 0.39–0.58 (Table 2). The ages from 23 concordant analyses ranging from 494 to 432 Ma (Fig. 10b) are grouped at 482 ± 4.1 Ma (6 grains, MSWD = 0.5, probability = 0.48), 458.1 ± 4 Ma (13 grains, MSWD = 3.8, probability = 0.05) and 433.2 ± 6.5 Ma (4 grains, MSWD = 2.1, probability = 0.15). The oldest group was mostly obtained from the core of the grain (Fig. 9b, G#16), whereas the youngest grain, 432 ± 13 Ma (MSWD = 1.05, probability = 0.31), was from a rim location (Fig. 9b, G#37-R; Table 2). Therefore, the youngest group age of ca. 433.2 ± 6.5 Ma is considered to be of the same age as the porphyritic granite magmatism (Fig. 10b).

Sample 12Hr1011

This sample is a clast in the cataclasite of the thrust. It is fine-grained porphyritic rhyolite with phenocrysts of quartz, K-feldspar, and plagioclase. The zircons from this sample are commonly needle-shaped or acicular, euhedral, transparent, and range from 50 to 150 µm in length along their major axis. Fewer cracks and apatite needles are included. Most grains show concentric oscillatory zoning in their CL images (e.g., Fig. 9c).

The Th/U ratio is 0.35–0.79 (Table 2). The ages from nine concordant analyses ranging from 524 to 424 Ma (Fig. 10c) are grouped at 457.1 ± 6.9 Ma (7 grains, MSWD = 4.2, probability = 0.04). The youngest age, 424 ± 18 Ma (one grain, MSWD = 0.56, probability = 0.46), was obtained from an oscillatory zoned rim (Fig. 9c, G#4-R), and the 484 ± 21 Ma and 454 ± 19 Ma ages were obtained from the core and rim of a single grain of zircon, respectively (Fig. 9c, G#6-C and G#6-R). The older (~524 Ma) and younger (~450–420 Ma) ages can, therefore, be interpreted as having been inherited which were extracted from crustal materials and magmatic porphyritic rhyolite.

Shortly, the U-Pb ages of 561–480, 492–432, and 524–424 Ma are yielded from zircons, which extracted from samples of the equigranular granite (#12Hr1601), porphyritic granite (#11Hr54), and porphyritic rhyolite (#12Hr1011), respectively. These U-Pb ages correspond to the Ediacaran to Late Silurian.

Discussion

Timing of the low-angle southward-dipping thrusting in the Züünharaa area

The low-angle southward-dipping thrust in this area clearly cuts the Haraa Group, the porphyritic granite–rhyolite, and the Ulaan Öndör Formation, with the exception the Örmögtei Formation. The granitic rocks had U-Pb ages ranging from 561 to 424 Ma (Figs. 2, 10) and coincide with the U-Pb ages of the Boroogol plutonic rock complex dated at 460–440 Ma (e.g., Kröner et al. 2007; Hou et al. 2010).

Magmatic zircon from the porphyritic granite–rhyolite clast in the cataclasite shows ages of ~450–420 Ma, and detrital zircons of ~450 Ma were reported from the Haraa Group (Bussien et al. 2011; Kröner et al. 2007; Kelty et al. 2008). This suggests that the low-angle south-dipping thrust occurred after the formation of the Haraa Group and the intrusion of the porphyritic granite–rhyolite into the Haraa Group. The tuff breccia of the Ulaan Öndör Formation, which contains abundant clasts of porphyritic granite–rhyolite and covers the Haraa Group and granitic rocks, is cut by the southward-dipping thrust. Yet, the clastic rocks Visean Örmögtei Formation were not affected by shearing. The thrusting can thus be considered to have occurred between the Silurian and the Visean.

Distribution of the low-angle southward-dipping thrust and regional correlation of the Haraa terrane

Several studies have reported low-angle faults similar to the low-angle southward-dipping thrust from the Haraa terrane (L.2–L.4, Fig. 1b, e.g., Tomur et al. 1994; Purevsuren and Narantsetseg 1998; Dejidmaa 2003; Tovuudorj 2003; Altanzul and Baasandolgor 2014). The meta-sandstone of the Cambrian–Early Ordovician Haraa Group thrusts over Middle–Late Ordovician porphyritic rhyolite at Zuunmod area, 30 km southeast of the study area (L.2, Fig. 1b; Altanzul and Baasandolgor 2014). The thrust plane strikes N 20° E and gently dips southward (Altanzul and Baasandolgor 2014). Dejidmaa (2003) reports that a sub-horizontal top-to-the-north fault cuts volcaniclastic rocks of the Devonian Ulaan Öndör Formations at Bayangol area, 40 km west of the study area (L.3, Fig. 1b). Kelty et al. (2008) suggested that the Carboniferous Örmögtei Formation is exposed as a tectonic window below the Cambrian–Early Ordovician Haraa Group in the Zaamar area, 150 km southwest of the study area (L.4, Fig. 1b). Although details of the ages of the faults still remain unknown, they are likely to originate in the Late Paleozoic based on the fact that it is generally hard to recognize low-angle faults in the Triassic/Jurassic granites distributed throughout the Haraa terrane (e.g., Tomurtogoo et al. 1998).

Late Paleozoic low-angle southward-dipping thrusts are also recognized in the SB belt in North Mongolia and West Transbaikal, northwest of HD belt (Fig. 1b) (e.g., Zorin 1999; Ruzhentsev et al. 2006, 2007, 2012; Buslov et al. 2009; Ryabinin et al. 2011; Zhimulev et al. 2011). Buslov et al. (2009) and Zhimulev et al. (2011) reported that the Neoproterozoic–Cambrian formations overthrust upon the Middle Ordovician Tunka granite and the Late Devonian–Early Carboniferous Sagan-Sair Formation with an E–W trending plane and dipping less than 30° southward in the Tunka area (L.4; Fig. 1b).

The minerals (such as amphibole, biotite and muscovite) at the bottom of the tectonic sheet of the Neoproterozoic-Cambrian formations (dated at an U-Pb age of 2–1.7 Ga; Zhimulev et al. 2010) yield an Ar–Ar age of 316–310 Ma, which correspond to the Late Carboniferous (e.g., Buslov et al. 2009; Ryabinin et al. 2011; Zhimulev et al. 2011). Ryabinin et al. (2011) and Zhimulev et al. (2011) interpreted this isotopic age as the timing of the low-angle southward-dipping thrusts.

Several examples of Late Paleozoic low-angle south-dipping thrusts can be seen in the Ul’zutui, Oldynda, Kydzhimit areas of the Eravna terrane, SB belt (L.5, Fig. 1b). The andesite dated at 310 Ma and felsite dated at 297 Ma occur as a series of tectonic sheets alternating with slices of Lower Paleozoic rocks along low-angle south-dipping plane in the Ul’zutui, Oldynda, and Kydzhimit areas of the Eravna terrane (Ruzhentsev et al. 2012). Another example of Late Paleozoic thrust is found in the Bagdarin area of the Ikat terrane (L.6, Fig. 1b). The Neoproterozoic-to-Middle Paleozoic formation thrusts onto the Late Devonian–Early Carboniferous Bagdarin Formation along low-angle south-dipping planes in this area (Ruzhentsev et al. 2007, 2012). The Usoi granitic rocks of the Angara-Vitim batholith (dated as 288 ± 2 Ma) cut both the hanging-wall and footwall of the low-angle south-dipping thrust in the Bagdarin area (Ruzhentsev et al. 2007; Mazukabsov et al. 2010). It can thus be concluded that the thrusts in the Ul’zutui, Oldynda, Kydzhimit, and Bagdarin areas were formed after the Early Carboniferous and before the Early Permian periods.

The above observations, i.e., L.1–L.6 (Fig. 1b), make it clear that the Late Paleozoic low-angle southward-dipping thrust can generally be recognized in both the HD and SB belts. It should be noted that Proterozoic–Early Paleozoic sedimentary rocks are intruded by Cambrian–Ordovician granitic rocks and unconformably overlain by Late Paleozoic sedimentary rocks in the SB belt (e.g., Parfenov et al. 2009; Buslov et al. 2013). In the Haraa terrane of the HD belt, however, Cambrian–Lower Ordovician sedimentary rocks are intruded by Cambrian–Silurian granitic rocks and are unconformably overlain by Devonian–Carboniferous volcaniclastic and clastic rocks (e.g., Tomurtogoo et al. 1998; Tomurtogoo 2012). Thus, the general Paleozoic stratigraphy of this terrane is substantially the same as that of the SB belt. It is generally accepted that the SB belt is a collage of blocks with granitic batholiths that were accreted to the Siberian craton during the Early Paleozoic (e.g., Parfenov et al. 2009), and which behaved as a part of the continent during the Paleozoic. Badarch (2005) expressed that the northwest of the Haraa terrane is the Precambrian continental basement rock beneath the Permian–Triassic superimposed volcanic–plutonic rocks, which is connected with the SC in the North Mongolia–West Transbaikal region (Fig. 1b). Together with the observations discussed above, it follows that the existence of the Late Paleozoic low-angle southward-dipping thrust and the stratigraphic similarity in both the Haraa terrane and the SB belt seems to demonstrate that the Haraa terrane can be correlated to the SB belt as a part of the continent.

Tectonic implication of the low-angle southward-dipping thrust

The Haraa terrane is limited by the NE-trending Late Mesozoic Yeroogol sinistral strike-slip fault system (e.g., Tseden et al. 1992; Kotlyar et al. 1998; Altanzul and Delgertsogt 2012; Altanzul and Baasandolgor 2014) at its southern end, as well as by its contact with the Asralt Hairhan terrane of HD belt (Fig. 1b). The Asralt Hairhan terrane is considered to have a metamorphic affinity to the Ulaanbaatar terrane (Tomurtogoo 2012; Gordienko et al. 2012). The Ulaanbaatar terrane is a late Devonian–Early Carboniferous accretionary complex (Kurihara et al. 2009). Several authors (e.g., Zorin 1999; Dorjsuren et al. 2006; Kurihara et al. 2009; Bussien et al. 2011; Gordienko et al. 2012; Takeuchi et al. 2012; Hara et al. 2013; Purevjav and Roser 2013; Ruppen et al. 2013; Tsukada et al. 2013) have been studied the lithology, stratigraphy, geological structure, geochemistry, and age of the rocks in the Paleozoic accretionary complexes of the HD belt. Detrital zircons in the sandstone of the accretionary complex of the Ulaanbaatar terrane mostly yield U-Pb ages of ~350–320 Ma (e.g., Kröner et al. 2007; Kelty et al. 2008; Bussien et al. 2011; Hara et al. 2013). The rocks of the accretionary complex are unconformably overlain by formations yielding Carboniferous brachiopods (e.g., Minjin et al. 2006; Kurihara et al. 2009; Takeuchi et al. 2012), and it can thus be inferred from the ages of the detrital zircon and overlying formation that the clastic rocks of the accretionary complex should be assigned to the Carboniferous. The sandstones of the accretionary complex and overlying formation in the Ulaanbaatar terrane (Fig. 5a, b, e.g., Bussien et al. 2011; Hara et al. 2013; Suzuki et al. 2012) are located in the “Basement uplift” and “Volcanic arc” fields in the Q–F–Lt diagram (Dickinson et al. 1983), as is the coeval sandstone of the Örmögtei Formation in the Haraa terrane. The sandstone of the Örmögtei Formation includes detrital zircon, most of which yields an U-Pb age of ~360–330 Ma (Kröner et al. 2007; Kelty et al. 2008). The similarities in detrital zircon ages and sandstone compositions in the Carboniferous formations of the Haraa and Ulaanbaatar/Asralt Hairhan terranes likely suggest that the sandstones were formed in adjacent areas. That is, there existed a continental affinity between the Haraa terrane and accretionary complexes of the Ulaanbaatar/Asralt Hairhan terranes, and they were probably located close together during the Carboniferous.

Several studies have described southeast-verging composite folds and northward-dipping thrusts in the accretionary complexes of the Asralt Hairhan and Ulaanbaatar terranes (L.8, Fig. 1b, e.g., Kurihara et al. 2009; Gordienko et al. 2012; Nakane et al. 2012; Suzuki et al. 2012; Takeuchi et al. 2012). These have been interpreted as showing that the folds and thrusts in these terranes were formed in relation to northward subduction of the previous oceanic plate beneath the continental crust (e.g., Kurihara et al. 2009; Gordienko et al. 2012; Takeuchi et al. 2012; Ruppen et al. 2013). Kurihara et al. (2009) proposed that the accretion process may have taken place between the late Devonian and Early Carboniferous, and Hara et al. (2013) inferred the accretion age of the clastic rocks as Early Carboniferous, based on the U-Pb age of detrital zircon in the Ulaanbaatar terrane. Early Permian mafic-to-felsic dikes cut the folded clastic rocks of the accretionary complex of the Ulaanbaatar terrane (Khishigsuren et al. 2009). Hence, the southeast-verging folds and northward-dipping thrusts are considered to have been formed before the Early Permian. In summary, Late Paleozoic southward-dipping thrusts are dominant in the Tunka, Ikat, and Eravna terranes of the SB belt and the Haraa terrane of the HD belt (L.1–L.7, Fig. 1b). In contrast, contemporaneous southeast-verging composite folds and northward-dipping thrusts have developed in the Asralt Hairhan and Ulaanbaatar terranes (L.8, Fig. 1b).

Simultaneously formed “doubly vergent (or bivergent; e.g., Willett et al. 1993; Storti et al. 2000; Naylor and Sinclair 2007; Mukherjee 2013b; Bose and Mukherjee 2015) asymmetric structures”, similar to those present in the HD and SB belt, have been illustrated in the Alps, the Andes, and in other locations. For instance, the Taranaki Fault in the northern part of New Zealand is considered to be a back thrust antithetic to the Hikurangi margin subduction thrusts. The Taranaki Fault has accommodated at least 12–15 km of dip-slip displacement since the middle Eocene (~40–43 Ma) (e.g., Stern et al. 2006; Nicol et al. 2007; Stagpoole and Nicol 2008).

The doubly verging character of the Eastern Alps architecture is evident from the predominant criss–cross reflection pattern at ~10 km depth in a 150–220 km interval of a ~300 km seismic section (Gebrande et al. 2006), along which giant crustal wedges have been upthrust since the Miocene (Pfiffner et al. 2000). The southward- and northward-dipping thrusts are exposed at the Inn Valley and the Valsugana-Agordo areas in the Eastern Alps, respectively (Slejko et al. 1989).

Another example can be seen in France and Spain. The doubly vergent asymmetric structure characterizes a ~150-km-wide surface expression of the Hercynian basement, thin-skinned fold-thrust belts shown by the 250-km-long deep seismic survey (ECORS Pyrenees profile; Choukfoune 1989) from the Aquitaine basin to the Ebro basin of the Pyrenees (Sinclair et al. 2005). This structure is considered to have formed in the Pyrenees as a result of northward subduction of the Iberian plate beneath the East European craton during Late Cretaceous to early Miocene times (Roest and Srivastava 1991; Sinclair et al. 2005).

The Andes are sustained by large doubly vergent thrust systems (West Andean and East Andean Thrust) close to latitude 21°S (e.g., McQuarrie et al. 2005; Armijo et al. 2015). This is a result of the protracted processes of doubly vergent crustal shortening and thickening in the more than 600-km-wide regions of the high Andes-Altiplano plateau, since 50 Ma has been compressed between the rigid Marginal Block and the South America Plate (Armijo et al. 2015).

As shown in the above examples, the doubly vergent asymmetric structure is common in the plate convergence fields, i.e., subduction and collision zones. The oceanward-verging folds and thrusts are developed in accretionary complexes, whereas the landward-verging thrusts are formed on the continental side under the Alpine-type compressional orogen (e.g., Beaumont et al. 1996; Poblet and Lisle 2011) such as Alps and Himalaya (Mukherjee et al. 2013, 2015).

In the HD and SB belts, the area of southeastwardly-verging composite folds and northward-dipping thrusts (Asralt Hairhan and Ulaanbaatar terranes) exposes accretionary complexes. In contrast, the areas dominated by southward-dipping thrusts (Tunka, Ikat, Eravna, and Haraa terranes) are assigned to a part of the continent with Siberian craton. These contrasting structures suggest that there is a doubly vergent asymmetric structure in the North Mongolia–West Transbaikal region. However, the Permian–Triassic volcanic–plutonic rock complexes largely exposed in the areas between the Haraa terrane of the HD belt and the terranes of the SB belt, and the Late Paleozoic structure is obscure (Fig. 1b). The volcanic–plutonic rock complexes are interpreted to be the result of subduction-related magmatic activity along the Paleozoic continental margin of the Siberian craton (e.g., Donskaya et al. 2013).

We show the schematic section of the doubly vergent asymmetric structure within the HD belt (Fig. 11) because of lack of Late Paleozoic structures in the Permian–Triassic superimposed volcanic–plutonic rock complexes (Fig. 1b). The Late Mesozoic Yeroogol sinistral strike-slip fault system marks the geological and structural boundary between the Haraa terrane and the Asralt Hairhan terrane of the HD belt (Figs. 1b, 11; e.g., Tseden et al. 1992; Kotlyar et al. 1998; Badarch et al. 2002; Tomurtogoo 2003, 2012). This fault cuts the Triassic/Jurassic granites which are distributed in both Haraa and Asralt Hairhan terrane (Tomurtogoo et al. 1998); displacement of fault is ~1 km in the Zuunmod area (L.2, Fig. 1b; Altanzul and Delgertsogt 2012). Altanzul and Baasandolgor (2014) inferred that the Mesozoic Yeroogol sinistral strike-slip fault cuts the low-angle south-dipping thrust at Zuunmod area (Fig. 11b). By contrast, Gordienko et al. (2012) mentioned that there were a few northward-dipping thrusts in south of the Zuunmod area. The facts with ~1 km displacement of the Mesozoic Yeroogol fault suggests that a symmetry of the doubly vergent asymmetric structure in the HD belt may possibly be found near the Zuunmod area (L.2, Figs. 1b, 11). It can be concluded that the doubly vergent asymmetric structure was formed before the Mesozoic Yeroogol sinistral strike-slip fault. Thus, to understand the geometry and formation time of the doubly vergent asymmetric structure in the HD belt, further detailed study is essential.

Fig. 11
figure 11

Schematic section of the Late Paleozoic doubly vergent asymmetric structure in the HD belt. Blue lines are southward-verging composite folds of accretionary complexes in the Ulaanbaatar and Asralt Hairhan terranes. Red lines are south- and northward-dipping thrusts. The vertical line is the Mesozoic Yeroogol sinistral strike-slip fault. This schematic section is compiled by this study with previously published geological maps (e.g., Kurihara et al. 2009; Gordienko et al. 2012; Takeuchi et al. 2012; Altanzul and Baasandolgor 2014)

Full size image

Finally, the Paleozoic continental margin (Tunka, Ikat, Eravna and Haraa terrane) and Paleozoic accretionary complexes (Ulaanbaatar/Asralt Hairhan terranes) exists in the SB and HD belts. They were located in proximity during the Carboniferous based on their similarities in geological background. The former area (Fig. 1b) corresponds to the area of “oceanward-verging folds and thrusts” and “landward-verging thrusts” (Fig. 11) of Alpine-type compressional orogen. The doubly vergent asymmetric structure in the former area was formed by the northward subduction process of the Mongol Okhotsk oceanic plate (e.g., Zorin 1999; Bussien et al. 2011; Donskaya et al. 2013), which was previously present between SC and NCP/TB (Fig. 1a) in Late Paleozoic period.