—The Hercynian mobile belts in Central Asia include the proper Hercynian and late Hercynian (Indo-Sinian) belts, whose formation is associated with the evolution of the South and Inner Mongolian basins with oceanic crust. Within the South Altai metamorphic belt (SAMB), rock complexes compose tectonic slivers of different ranks. At the early stages, their metamorphic alteration occurred under conditions of the high-temperature subfacies of the amphibolite and, in places, granulite facies. Structurally, the band of the outcrop of these complexes is confined to the Caledonian North Asian continental margin and stretches along the southern slope of the Gobi–Mongolian–Chinese Altay Mountains from southeast to northwest (East Kazakhstan), where they occur in the Irtysh strike-slip zone. We assign these complexes to the Hercynian SAMB running for more than 1500 km. The latter comprises poly- and monometamorphic complexes. Late metamorphic granitoids of the Tseel tectonic sliver (Gobi Altay) in the southeast of the SAMB have been dated at 374 ± 2 and 360 ± 5 Ma. The previous data and these results show that the early (~390–385 Ma) low-pressure and late (375–360 Ma) high-pressure metamorphism proceeded almost along the entire belt. The interval between them was a short tectonic lull. These processes took place during the closure of a Tethyan basin of the South Mongolian Ocean (Paleo-Tethys I). The spatial position of the SAMB was controlled by the structural asymmetry of the basin, with an active continental margin at its northern edge and a passive one at the southern edge (in the present-day coordinates).

Orogenic complexes of Central Asia are characterized by two main types of tectonic structures: mosaical and linear (Mossakovskii et al., 1993; Didenko et al., 1994). Structures of the first type are mainly Baikalides and Caledonides, while structures of the second one are the Hercynian mobile belts of the western Altai–Sayan region and southern Mongolia (Fig. 1). Along their boundary with the Caledonides, the Gobi Altay zone, considered to be the marginal part of the Caledonian paleocontinent, is located (Ruzhentsev et al., 1990; Ruzhentsev and Pospelov, 1992).

Tectonically, formation of the South Altai metamorphic belt was controlled by its position along the junction zone of the Caledonides and Hercynides (Fig. 1). The paper summarizes the results of structural, geological, and geochronological studies, allowing us to compile a model of formation of the Hercynian mono- and polymetamorphic complexes. The revealed features of the formation and evolution of these belts permit determination of the geodynamic setting of regional metamorphism in linear accretion–collisional structures in the central segment of the Central Asian fold belt.

In the Gobi Altay zone along the southern slope of the Gobi–Mongolian–Chinese Altay Mountains, as well as in East Kazakhstan (in the Irtysh strike-slip zone), rocks comprise tectonic slivers of different ranks, assigned to the single South Altai metamorphic belt (SAMB). These tectonic slivers have a width from hundreds of meters to 15–20 km and a length from a few to 50–60 km. In the north they are bounded by a system of ductile dislocations parallel to the northwest-striking structures of Paleozoic greenschist units, and in the south, they are cut by northwest- to sublatitudinally striking mylonite zones associated with deep faults (Irtysh, Bulgan, and Trans-Altai). The latter separate the SAMB from the Hercynian island-arc and oceanic complexes (Fig. 1). Along the entire SAMB, the tectonic slivers contain biotite and garnet–biotite gneisses with staurolite, kyanite, andalusite, and sillimanite as well as biotite–hornblende gneisses, amphibole schists, and amphibolites formed by metamorphism of sedimentary and volcanic precursors. These rocks are intruded by syn- and postmetamorphic granitoids and metabasite dikes. Many slivers demonstrate evidence of polymetamorphic processes, initially revealed only by structural geological methods and tentatively attributed to the Precambrian (Kozakov, 1986). Later these rocks were dated by zircon U–Pb analysis (Bibikova et al., 1992)1, providing grounds for their correlation with the Hercynides, and were examined applying petrological methods (Kozakov et al., 2002, 2011; Sukhorukov, 2007; Polyansky et al., 2011; Sukhorukov et al., 2016). Within the belt, the slivers of the metamorphic rocks are separated by shear zones, oriented, as a rule, in accordance with the main northwestern (less often, sublatitudinal) strike of the belt.

PT conditions of metamorphism

The tectonic slivers of the SAMB bear indications of zonal regional metamorphism of the kyanite–sillimanite facies of the M2 series. Analysis of staurolite and kyanite parageneses in the Mongolian Altay yielded temperatures and pressures of 560–630 °C and 5.5–7.5 kbar, respectively (Kozakov et al., 2002, 2011). These parageneses are also widespread in the tectonic slivers of the Chinese Altay and the Irtysh strike-slip zone of East Kazakhstan (Fig. 1), where the garnet–kyanite–biotite gneisses of the Predgornoe sheet gave close PT values of metamorphic conditions: 580–600 °C and 5.8–6.2 kbar, respectively (Kozakov et al., 2011).

The early metamorphic parageneses include the relict parageneses of the andalusite–sillimanite facies, reaching conditions of the granulite facies (M1); this event was accompanied by intense migmatization. Rocks with the early metamorphic parageneses occur only in the form of relics in boudins or tectonic lenses. The Tsogt sliver of the Gobi Altay (Fig. 2) contains granulites formed at 870 °C and 5.7 kbar, as was determined for the plagioclase + orthopyroxene + clinopyroxene + biotite + quartz paragenesis (Kozakov et al., 2002). The M1 metamorphic event in the Bodonchin sliver took place at lower pressures and temperatures, ranging from 520 to 560 °C and from 3.0 to 3.6 kbar, respectively. The staurolite and kyanite parageneses of the late metamorphic episode (M2) most widely occurred in the Mongolian Altay; the PT values range from 560 and 630 °C and from 5.5 to 7.5 kbar, respectively (Kozakov et al., 2011). In the Bulgan tectonic sliver, hypersthene relics occur among the diopside-bearing gneisses composing the central parts of the boudins, hosted by the gneisses with the high-pressure metamorphic paragenesis of kyanite + garnet + staurolite + muscovite + biotite + plagioclase + quartz (Kozakov, 1986). This suggests formation of these rocks by metamorphic alteration of previous highly metamorphosed precursors.

The structure of most of the tectonic slivers in the belt was finally formed by the late episode of high-pressure metamorphism (M2) and most of the internal deformations associated with it. Notably, the overlapping metamorphism in many cases “erases” the early (M1) metamorphic parageneses.


The early (M1) episode of metamorphism has been dated by analysis of synmetamorphic amphibole monzodiorites at 385 ± 2 Ma (Bibikova et al., 1992); the rock intrudes into migmatized hornblende gneisses with relics of hypersthene gneisses in the Tsogt tectonic sliver of the SAMB and, in turn, is cut by the Gashunnur complex dikes under conditions of amphibolite metamorphism (M2). The age of the later, high-pressure metamorphic episode (M2) has been determined using zircons from synmetamorphic layered gabbro and trondhjemites of the Ekhnii-us massif (Bodonchin tectonic sliver) at 371 ± 2 and 365 ± 4 Ma, respectively (Bibikova et al., 1992). The gabbro-diorites of the massif contain xenoliths of folded migmatized gneisses associated with the earlier episode of metamorphism. Similar age estimates for the early and late metamorphic events (384 ± 2 and 358 ± 6 Ma) have been obtained on metamorphogenic zircons from the migmatized gneisses containing relics of hypersthene (Tsogt sliver, Gegetin-Gol River valley) (Bibikova et al., 1992; Kozakov et al., 2002).

Estimates of the early limit of accumulation of metasiliciclastics in the Bodonchin tectonic sliver at 458 ± 4.5 and 550–450 Ma have been obtained by dating detrital zircons, while the early episode of 385 ± 5 Ma old metamorphism (Kozakov et al., 2009; Jiang et al., 2012) is accepted as the sedimentation termination limit. That is, the time span of their accumulation did not exceed 60–70 Myr and corresponds to the Late Ordovician–Early Devonian.

In the Chinese Altay, the SAMB tectonic slivers (Fig. 2) contain gneisses with staurolite, kyanite, and garnet along with andalusite-, cordierite-, and sillimanite-bearing varieties (Windley et al., 2002; Zheng et al., 2007). There the granitoids terminating the late metamorphic episode have been dated at 373 ± 2 Ma (Kozakov et al., 2011) (Fig. 3). In the northwestern part of the Chinese Altay, the 462 ± 10 Ma old granitoids, metamorphosed together with the host rocks under conditions of the amphibolite facies, contain meta-morphogenic zircon dated at 400–370 Ma (Wang et al., 2006). In East Kazakhstan, the northwestern continuation of the SAMB (Fig. 1) hosts sheared 226 ± 9 Ma old granitoids occurring in strike-slip zones bordering the tectonic slivers (Kozakov et al., 2011). At that, the age of metamorphic zircon from ultrametagenic garnet-bearing granites of the Kurchum–Kal’dzhir sliver is 362 ± 5 Ma (Kozakov et al., 2011). It should be noted that this sliver was metamorphosed under conditions of the andalusite–sillimanite facies and no early episode of metamorphism is recognized, which distinguishes it from the rest of the SAMB slivers.

In general, within the eastern part of the SAMB, the rocks had formed by the beginning of the early Carboniferous and later underwent no alteration under conditions of high-temperature regional metamorphism. For instance, the Permian granitoids are not affected by strike-slip zones. Yet, in the northwestern continuation of the belt in the Chinese Altay and East Kazakhstan, where the metamorphic processes ceased in the Devonian, the Permian granitoids are known to experience mainly low-temperature metamorphic transformations.

The Gashunnur metabasite complex

The dikes of the Gashunnur complex intruded between the early and late metamorphic episodes at ~380–375 Ma. To different extents they occur in many tectonic slivers of the SAMB. They are most abundant in the Tseel sliver (Fig. 4), where they form metabasite dike fields and swarms (Fig. 5 a, b). In the area of Tseel soum, their geochemical affinity assumes similarity to basalts of midocean ridges or oceanic plateaus. In the southeastern part of the Tseel sliver, their compositions shift to that of volcanic-arc basalts. Samarium–neodymium isotope studies corroborate these differences in their sources of melt (Kozakov et al., 2019).

Structures of the Tseel tectonic sliver

The chronology of structural evolution of the Tseel tectonic sliver is based upon correlation with the stages of development of the Bodonchin and Tsogt tectonic sheet structures; the last two are located in the northwestern and southeastern continuations of the SAMB (Fig. 2). Some LA- ICP-MS zircon ages, ranging from 550–460 to 430–350 Ma (Burenjargal et al., 2014; Hanžl et al., 2016), have been obtained in the Tseel sliver, but they do not provide a ground for correlation of the dated objects with stages of the structural and metamorphic evolution of the SAMB. The folded structures of the early metamorphic episode (M1) in the Tseel sliver and, as a rule, in other tectonic sheets had initially a submeridional strike (modern-day coordinates); that is, it did not coincide with the main northwestern and latitudinal attitude of the SAMB tectonic slivers (Kozakov, 1986; Kozakov et al., 2007, 2011). This is seen in the orientation of relict fold hinges and mineral lineation (Fig. 4). Superposition of the regional late metamorphism (M2) was accompanied, at an early stage, by formation of low-angle structures: recumbent isoclinal folds and newly formed gently dipping schistosity and metamorphic banding parallel to their axial planes (Fig. 6). As a rule, the superposed metamorphism was not accompanied by migmatization, though sometimes feldspathization is recorded. The structures and parageneses of the early metamorphism (M1) are preserved only as relics (Fig. 5,c), which are bounded by shear zones and, in most cases, obscured. The general structures of the Tseel sliver are straight isoclinal folds of sublatitudinal strike produced by the late metamorphic episode (M2) (Fig. 2) as well as regionally developed schistosity parallel to their axial planes (Kozakov et al., 2007, 2011). In the Tsogt, Barlagin, and Bodonchin slivers, these strike northwest; that is, they mainly follow the dominant structure of the SAMB (Fig. 2). The chronology of formation of the Tseel sliver structures has not been revealed.

To estimate the age of the M2 metamorphism, a synmetamorphic quartz diorite (sample 8136) has been collected (Fig. 4). The synmetamorphic nature of this 374 ± 2 Ma old quartz diorite is determined by the presence of crystallization schistosity (on biotite, less often with hornblende), parallel to the axial planes of the main isoclinal folds formed during the late regional metamorphic stage in the host rocks (biotite gneisses with amphibolites). The occurrence of crystallization schistosity implies that by the time of its formation the quartz diorites had already crystallized together with dikes of the late-phase plagiogranite. The low-angle structures (crystallization schistosity and recumbent folds of the first stage of the M2 metamorphism), which regionally affected the dikes of the Gashunnur complex and migmatites of the early metamorphism (M1), are not pronounced in the quartz diorite (Fig. 6). Hence, their intrusion took place after the formation of the low-angle structures of the M1 stage, but prior to the formation of upright folds and formation of the regional schistosity of the late metamorphic episode (M2). The northward vergence of the folds and schistosity in the massif in the Tseel sliver is associated with thrusting which occurred during juxtaposition of the metamorphic rocks against the Early Paleozoic complexes of the southern margin of the Caledonian paleocontinent (Fig. 4). These relations essentially coincide with those found in the Ekhnii-us massif (371 ± 2 Ma) of the Bodonchin sheet: The intrusion took place between the first and second stages of the M2 metamorphism. In the Bodonchin, Barlagin, and Tsogt tectonic sheets, their zones of junction with the northerly framing complexes have a subvertical plunging, to a lesser extent complicated by thrusts. The intrusive rocks associated with this metamorphism are sheared in different degrees, and their schistosity is concordant with the structures of the host metamorphic rocks.

The late age limit of the M2 metamorphism is determined by dating of massive postmetamorphic subalkaline granite (sample 8200), collected in the southeast of the Tseel sliver. Their host rocks comprise abundant migmatite of the early metamorphic stage M1, anatectic granitoids, and the Gas-hunnur metabasites which intrude into them and form dike swarms (Fig. 4). The granites, in fact, crosscut the Gashunnur metabasite dike field; hence, they postdate the M2 metamorphism. Throughout the massif, they demonstrate no signs of schistosity or thermal impact on the well-pronounced oscillatory zoning of magmatic zircons. However, the amphibolite-facies metamorphism is recorded in the host rocks directly along the entire contact. The occurrence of the sheeted granite body in an open fold (Fig. 4) may be associated with later postmetamorphic shear deformations that took place during assemblage of the Tseel tectonic plate with the Hercynian complexes of the southern framing.

Analytical methods

The zircon was separated following the standard technique using heavy liquids. The zircon crystals (or their fragments) selected for U–Pb analysis were subjected to multistage removal of surface contaminants in alcohol, acetone, and 1 M HNO3. After each step, the grains were rinsed with ultrapure water. The chemical digestion of the zircon and separation of U and Pb were carried out following the modified technique of Krogh (1973). Isotope analyses were performed employing a TRITON TI multicollector mass spectrometer both in static mode and using an ion counter. The samples were spiked with a 235U–202Pb tracer. Accuracy of the U and Pb determination was 0.5%. The blank contamination did not exceed 15 pg Pb and 1 pg U. The acquired data were processed using PbDAT (Ludwig, 1991) and ISOPLOT (Ludwig, 2003) macros. The ages were calculated applying generally accepted decay constants of uranium (Steiger and Jäger, 1977). Correction for common lead was done in accordance with model values (Stacey and Kramers, 1975). All the errors are given at the 2σ level.

The accessory zircon from the quartz diorite (sample 8136) is observed as translucent, colorless or yellowish euhedral and, less often, subhedral crystals. The sample contains short-prismatic, prismatic, and long-prismatic (up to asicular) crystals. Their size varies from 50 to 400 μm, with aspect from 1.0 to 7.0. The faceting is represented by a combination of a {100} prism and {101}, {201} bipyramids (Fig. 7 a, I–IV).

The internal structure of the grains was examined using an optical microscope and SEM with a cathodoluminescence detector. The zircon has a clearly pronounced fine-zoned internal structure (Fig. 7,a, V–VIII); it contains clusters of tiny gas–liquid and opaque mineral inclusions. Individual crystals contain inherited cores (Fig. 7 a, II).

The U–Pb analyses were carried out on four microspecimens of most clear and euhedral zircons from 50–100 and 150–200 μm fractions (Table 1). The analyzed zircon yielded results with an insignificant (1.8–4.8%) positive discordance. However, the position of figurative points on the Concordia diagram (Fig. 8,a) indicates that the age discordance is caused by both modern Pb loss (Nos. 1–3, Table 1) and presence of inherited lead (No. 4, Table 1). The average 206Pb/207Pb age of the analyzed zircon (Nos. 1–3) of 374 ± 2 Ma (MSWD = 1.8) may be used as the best estimate of its crystallization.

In the granitoids, sample 8200, the zircons are euhedral to subhedral crystals of prismatic tracht with aspect of 1.8–3.6. The facets are prisms {100}, {110} and bipyramids {101}, {111}, {401} (Fig. 7,b, I–IV). The crystals are transparent to translucent, light yellow and colorless. They demonstrate a fine magmatic zoning (Fig. 7 b, V–VIII) and contain a large number of gas–liquid and mineral pulverulent and acicular inclusions.

The U–Pb dating was carried out on the most transparent and euhedral zircons (three microspecimens from size fractions 50–75 and 75–100 μm). Three results aligned to a regression line with an upper intersection with the Concordia at 360 ± 5 Ma (MSWD = 1.0) (Fig. 8,b, Table 1). The zircon morphology and structure indicate its magmatic origin; thus, the obtained age can be considered the time of formation of the granitoids.

The obtained dates of the syn- and postmetamorphic granitoids (374 ± 2 and 360 ± 5 Ma) imply that the late episode of metamorphism (M2) in the Tseel tectonic sliver of the SAMB lasted from ~375 to 360 Ma. The ages of the synmetamorphic layered gabbro of the early phase and the late-phase trondhjemite of the Ekhnii-us massif in the Bodonchin tectonic sheet embrace almost the same time span. Similar age estimates of the late metamorphic episode were determined throughout the SAMB: in the Chinese Al-tay (373 ± 2 Ma), the Tsogt sliver of the Gobi Altay (358 ± 6 and 364 ± 4 Ma (Kröner et al., 2010)), and in the Kurchum–Kal’dzhir sliver (362 ± 5 Ma) in East Kazakhstan. In the Kurchum–KaTdzhir sliver, the early metamorphism, which affected the rest of the SAMB tectonic slivers, is not recorded. However, amphibolite-facies metamorphism of similar age took place there under conditions of reduced pressure, contrasting with that in other slivers, including the adjacent tectonic sheets of the Irtysh strike-slip zone of East Kazakhstan. It can be assumed that during this period the Kurchum–Kal’dzhir sliver was spatially separated from the main zone of formation of the SAMB. In the modern structure, it is separated by a thick zone of greenschist blastomy- lonites from the adjacent Predgornoe, Chechek, and Marka-kol’ sheets; the last three bear parageneses of the amphibolite facies of the kyanite-sillimanite facies series.

The effect of the early metamorphic episode (M1) is known in many tectonic slivers almost along the entire SAMB. In the central part of the Tsogt tectonic sheet, it was found to be ~385 Ma old (Bibikova et al., 1992), while in its southeastern extension, it was dated at 397 ± 3 and 396 ± 3 Ma (Demoux et al., 2009; Kröner et al., 2010). In the Tseel sliver, the gneissose granites yielded a close age of 385 ± 7 Ma (Burenjargal et al., 2014). In the Bodonchin sheet, the early and late episodes of the metamorphism are evidenced by the presence of xenoliths of migmatized 371 ± 2 Ma old amphibolites within the Ekhnii-us gabbro-diorites of the massif as well as by the presence of metamorphic zircons in paragneiss which are dated at 389 ± 3, 394 ± 1, and 374 ± 1 Ma (Jiang et al., 2012). The latter value almost concurs with the estimated age of the late metamorphic episode (M2). In the Sogra and Bukhtarma tectonic slivers in East Kazakhstan at the northwestern continuation of the SAMB, the gneisses and schists of the high-pressure amphibolite facies contain the earlier relict parageneses of the amphibolite facies of the andalusite–sillimanite facies series (Ermolov et al., 1984).

It should be emphasized that the metamorphic transformations of the recognized episodes are not evolutionary stages of a single metamorphic cycle which occurred during the ~390–360 Ma span, but rather events separated by a relatively short lull (~380–385 Ma); this suggests a change of the tectonic settings. First of all, this is evidenced by the Gashunnur basic dikes, intruding into the granites that terminate the early metamorphic episode (M1) processes. The granites form massifs and crosscutting veins in the migmatite; later, both were involved in structural–metamorphic transformations during the second episode of metamorphism (M2) under conditions of the amphibolite facies (Kozakov et al., 2007, 2019).

Considering the possible tectonic setting of formation of the SAMB, note that by the end of the Cambrian–beginning of the Ordovician, the early Caledonian paleocontinent had assembled, and a passive margin had formed at its southern edge (Ruzhentsev and Pospelov, 1992). In the early–middle Paleozoic, the passive margin transformed into an active one, and an accretionary wedge formed. The age spectrum of detrital zircons assumes a significant contribution of early Paleozoic sources in the provenance area, but also indicates a subordinate input from older, early and late Precambrian complexes (Jiang et al., 2011). In the Tseel tectonic sliver, detrital zircon cores from the pelitic gneisses gave ages ranging from 455 to 2600 Ma, dominated by the ~490–520 Ma population (Jiang et al., 2011; Burenjargal et al., 2014). The isotopic Pb and Nd characteristics of the middle and late Paleozoic granitoids of the Mongolian and Gobi Altay also demonstrate heterogeneity of their sources, which supposedly were metasiliciclastics of the mentioned accretionary wedge. They might have formed as products of erosion of rocks with Pb and Nd isotope affinity of island-arc complexes of the Lake Zone as well as the early and late Precambrian terranes that framed the south of the Siberian Platform and were sources of metamorphic rocks with old crustal isotopic characteristics (Savatenkov et al., 2020).

From the end of the Early Devonian and in the Middle Devonian, a system of island arcs and backarc troughs formed within the South Mongolian paleoocean, which is demonstrated by rapid growth of the juvenile crust (Yarmolyuk et al., 2007). By the Middle Devonian, spreading had ceased in all the oceanic basins, and activity in the extended subduction zones along the Siberian continent and Kazakhstan had led to the beginning of their effective closure.

The results of the studies allow us to propose the following sequence of formation of the SAMB mono- and polymetamorphic complexes (Fig. 9). In the Early Devonian, the convergence processes began in oceanic basins at the Caledonian paleocontinent boundaries, where the tectonic setting transformed from passive- to active-margin mode. This process is associated with the early episode of regional metamorphism (M1) during the ~390–380 Ma time span. The high-temperature and low-pressure conditions of this event suggest a low-angle subduction of a hot oceanic plate with the spreading axis proximal to the margin of the Caledonian Siberian paleocontinent (Fig. 9,a). During the subduction, the spreading center approached the newly formed active margin of the Caledonian paleocontinent and was subsequently subducted beneath it (Fig. 9 b). Its passage under the edge of the continent might have triggered a rifting process within the last subduction process, causing short-term stabilization and intrusion of the postmetamorphic (in relation to the early episode of metamorphism M1) granitoids. The formation of the Gashunnur dike swarms might be related to subduction of the spreading axis below the continent edge.

The Late Devonian (370–360 Ma) episode of high-pressure metamorphism (M2) correlates with the continuation of the middle Paleozoic accretion of island arcs in southern Mongolia (Yarmolyuk et al., 2007). It is associated with the formation of the tectonic slivers of different ranks (Fig. 2) as well as the formation of their sheeted internal structure. The resumption of the subduction process led to the regional metamorphism at increased pressure, associated with alteration of the early metamorphites along with the late M1 granitoids and dikes of the Gashunnur complex.

By the Middle Devonian, the spreading had ceased in all the oceanic basins, and as a result of activity in the extended subduction zones along the Siberian paleocontinent and Kazakhstan, their effective closure had begun. Presumably the subduction zones plunged down the newly formed margin of the North Asian paleocontinent and down a system of intraoceanic island arcs, displacing them toward the ocean (Ruzhentsev and Pospelov, 1992). The development of the active continental margin here began at ~400 Ma, but it had ended by the beginning of the early Carboniferous. The Ob’–Zaisan and South Mongolian oceans turned into residual oceanic basins (Ruzhentsev et al., 1990), and at the end of the Middle–beginning of the late Carboniferous, the Ob’– Zaisan and South Mongolian oceans closed completely.

In the Mongolian Altay, the early M1 metamorphism generally affected a relatively narrow strip of the southern parts of tectonic slivers. The structural and metamorphic transformations associated with the M2 metamorphism affected the previously metamorphosed rocks, but they are also pronounced in rocks with no parageneses or structures associated with the earlier metamorphism. That is, these formations were juxtaposed before the onset of the high-pressure metamorphism (M2). The M2 metamorphic processes altered both the rocks reaching the level of the high-temperature amphibolite and granulite facies and the rocks of lower degrees of M1 metamorphism.

A change of the thermodynamic regime of metamorphism from the low- to the high-pressure one in the “counterclockwise” trend may be associated with stabilization and, accordingly, cooling of the formed metamorphic rocks. During the continuation of the accretion–collisional processes and the recommencement of the subduction, they moved off the edge of the Caledonian paleocontinent, possibly providing a swap of the thermodynamic regime (Likhanov, 2020), assumingly due to a change of the subduction zone pitch.

Notably, the relict structures bearing signs of the early metamorphism M1, as well as the Gashunnur complex dikes, initially, as a rule, had a submeredional strike, nonconformable with the main structures formed by the late metamorphism M2 (Fig. 2). At the initial stage of the latter (after the intrusion of the Gashunnur dikes), the regional leveling of the early metamorphic structures occurred, producing the recumbent folds and low-angle schistosity (Fig. 6,a). This was probably caused by thrusting in the Earth’s crust deep levels, under conditions corresponding to the amphibolite and granulite facies. After that, the regional straight isoclinal folds and crystallization schistosity formed almost in the entire volume of the tectonic slivers of the SAMB, which is recorded in its modern structure (Fig. 6 b). Remarkably, these structural-metamorphic transformations did not affect the rocks of its framing. In particular, the Vendian–Cambrian rocks contacting the metamorphic rocks of the Tseel sliver display no signs of structural–metamorphic alteration (Markova and Fedorova, 1971). Presumably, the modern position of the SAMB tectonic slivers in the regional structure was caused by later shear deformations. The latter produced mylonitization or foliation at greenschist, sometimes up to the amphibolite-facies conditions. An exception is the tectonic slivers in the northwestern Chinese Altay, where Paleozoic (462 ± 10 Myr) granitoids are found to be metamorphosed together with the host rocks under conditions of the amphibolite facies; these granitoids bear metamorphic zircons dated at 400–370 Ma (Wang et al., 2006). That is, the SAMB regional metamorphism in this case affected the early Paleozoic units as well. In general, juxtaposition of the metamorphic rocks with the Vendian–early Paleozoic complexes of the northern framing occurred along postmetamorphic thrusts, and then both were dislocated by strike-slip deformations.

The obtained ages of syn- and postmetamorphic granitoids (374 ± 2 and 360 ± 5 Ma, respectively) allow us to estimate the duration of the late episode of high-pressure metamorphism (M2) in the Tseel tectonic sliver of the SAMB at ~15 Ma, from ~375 to 360 Ma. The dates are almost contemporaneous with the metamorphism which affected tectonic slivers of the Gobi and Mongolian Altay southeasterly and northwesterly of the Tseel sheet. Together the obtained and published geochronological and geological data provide grounds for correlating these metamorphic processes with those identified in the Chinese Altay and East Kazakhstan (Irtysh strike-slip zone) and associating them with evolution of a single Hercynian metamorphic belt.

Structural, geological, geochronological, and petrological data evidence the earlier episode of high-temperature, moderately low-pressure 395–385 Ma old metamorphism in the tectonic slivers of the Mongolian and Gobi Altay. A similar sequence of metamorphic processes has been recognized within the Chinese Altay and East Kazakhstan, yet no reliable geochronological data corroborating the early episode of metamorphism have been obtained.

The SAMB formation reflects the main stages of the Devonian evolution of the Caledonian paleocontinent margin. Subduction of a young oceanic crust in proximity to the spreading axis caused an early episode of the M1 metamorphism in the ~390–380 Ma interval. Dipping of the spreading ridge down the active margin of the Caledonian paleocontinent and opening of an asthenospheric window initiated a lull in the subduction activity and terminated the early meta-morphic episode during the ~380–375 Ma span. The late episode of metamorphism M2 lasted from ~370 to 360 Ma, when the accretion–collisional process resumed and island-arc complexes of the South Mongolian and Ob’–Zaisan paleooceanic basins were juxtaposed to the Altai margin of the Siberian and Kazakhstan paleocontinents. Presumably, the structural rearrangement during the transition to the late episode of metamorphism is associated with the temporal stabilization and related cooling of the metamorphic rocks formed to this time. The resumption of the accretion–collisional process and onset of the late metamorphic episode may have occurred during the amalgamation and common deformation of the rocks of different rheologies: the consolidated meta-morphic rocks and rocks of siliciclastic–volcanic complexes that previously experienced no metamorphic alteration.

In general, the SAMB formation is associated with the evolution of the Hercynian South Mongolian paleooceanic basin; its multistage development affected the Caledonian paleocontinent margins. However, the modern position of the tectonic slivers of the SAMB was determined by rearrangement by the later strike-slip displacement separated in time from the regional metamorphism, as it has been demonstrated in the Siberian Platform framing (Metelkin, 2012).

The authors are grateful to the reviewers’ constructive suggestions and comments, which allowed improvement of the article.

This work was financially supported by the Russian Foundation for Basic Research (project No. 20-05-00297).

In (Hanžl et al., 2016), it is erroneously stated that the age from (Bibikova et al., 1992) was obtained by Pb–Pb evaporation. This method was not used in the authors’ studies; zircon U–Pb dating (ID TIMS) was applied; the application of other methods is mentioned in the text.