Thermochronology of the Kalba–Narym Batholith and the Irtysh Shear Zone (Altai Accretion–Collision System): Geodynamic Implications

Abstract

The granitoids of the Kalba–Narym batholith and the Irtysh shear zone (ISZ) are among the main geological features of the late Paleozoic Altai accretion–collision system (AACS) in Eastern Kazakhstan. Traditionally, it is believed that late Paleozoic strike-slip faults played a pivotal role at all stages of the development of the AACS, they were supposed to control deformation, magmatism, and ore deposits. This work is devoted to solving the problem of the tectonic evolution of the AACS based on the reconstruction of the thermal history of granitoids of the Kalba–Narym batholith in connection with the Chechek metamorphic dome structure, which is one of the highly metamorphosed blocks mapped within the ISZ. The new geological and geochronological data presented in this work allowed us to establish the sequence of formation of the Kalba–Narym granitoid batholith and link it with the evolution of the Irtysh shear zone (ISZ). It was revealed that in the late Carboniferous–early Permian (312–289 Ma), during the NE–SW compression, the Irtysh shear zone formed as a gently dipping thrust system into which gabbro of the Surov massif intruded. The combined manifestation of magmatic and tectonic processes caused the formation of tectonic mélange with cataclastic gabbro and metamorphic rocks of the Chechek metamorphic dome structure (312–289 Ma). Compression caused the formation of a cover-thrust structure. The thickening of the crust under the probable thermal action of the Tarim plume led to the formation of the early Permian Kalba–Narym batholith (297–284 Ma) within the Kalba–Narym terrane. Denudation of the orogen occurred before the Early Triassic (280–229 Ma). In this way the sequence of formation of the Kalba–Narym batholith and the ISZ is consistent with the concepts of the stages of plume-lithosphere interaction within the AACS under the influence of the late Carboniferous–early Permian Tarim igneous province, but in the cover-thrust tectonic setting.

1. Introduction

The late Paleozoic Altai accretion–collision system (AACS) is traditionally distinguished on the territory of Eastern Kazakhstan [1,2,3,4,5], which is part of the large Central Asian fold belt (CAFB). CAFB resulted from the accretion–collision interactions of the Paleo-Asian Ocean (PAO) plate with the Siberian and Kazakh paleocontinents. The belt contains fold zones of different ages, formed during the Vendian–Paleozoic by successive accretion and collision of island arcs, microcontinents, and oceanic plateaus to the Siberian continent [1,6,7,8,9,10,11,12,13,14]. According to [15,16,17,18,19,20,21,22], a single Vendian–Paleozoic subduction zone developed in the PAO over which the Kipchak arc formed. During the Paleozoic, due to the drift and rotation of the Siberian and East European continents, deformations of the arc occurred, manifested in the formation of oroclinal bends and a large-amplitude strike-slip system. The most important episodes of strike-slip movements of the terranes are the late Carboniferous and then late Permian. According to this model, fragments of the Kipchak arc, originally framing the Siberian and East European continents, were combined in the CAFB (Altaids, according to [15]) since the late Paleozoic.

The late Precambrian–Paleozoic collisional and accretionary orogenies have been distinguished in the geodynamic zoning of the CAFB (Figure 1) [23,24,25].

Figure 1. Geodynamic scheme of the Central Asian folded belt and the position of the studied object (Figure 2) (after [24,25]).

Figure 1. Geodynamic scheme of the Central Asian folded belt and the position of the studied object (Figure 2) (after [24,25]).

A characteristic feature of collisional orogens of the CAFB is the presence of Precambrian microcontinents of the Gondwana group among the units involved in the collision. At the same time, a characteristic feature of accretion orogens of the CAFB is the absence of microcontinents of the Gondwana group among the units involved in the collision [23,24,25]. If the two types of orogens mentioned above have been relatively preserved near the Siberian craton (Altai-Sayan region, Tuva, Mongolia, Transbaikalia), then in the rest of the CAFB their position and relationship are largely disrupted by late Paleozoic thrusting and strike-slip faulting. To the greatest extent, the primary structure is distorted by late fault tectonics within the AACS (Figure 1).

Several large fault and/or shear zones (depending on the depth of the erosion section) of the northwestern strike are manifested in the AACS. The largest among them are Zharma, Chara, Irtysh, Northeastern, and others. They limit large (with a width of many tens to hundreds of km) terranes, such as Chingiz–Tarbagatai, Zharma–Saur, Zapadno–Kalbin, Kalba–Narym, Rudny–Altai, and Gorny–Altai (Figure 1). These fault zones controlled tectonic deformations, metamorphism, magmatism, and ore deposits in the AACS [1,2,4,13,14,26,27,28,29,30,31,32,33,34,35,36,37,38,39]. According to [13], in the AACS in the late Paleozoic–early Mesozoic, accretion–collision processes manifested themselves as structure-forming factors, while the Tarim and Siberian plumes played the role of energy sources [40,41,42,43,44,45,46,47,48,49,50,51].

If the large late Paleozoic strike-slip systems of the northern part of the Central Asian Basin are well studied and characterized [4,24,27,28,52,53], then the accompanying and (or) preceding covering structures have still been poorly identified [24,52,54]. In the case of a high degree of denudation, vertical root zones of cover and shear structures are preserved on the erosion section affecting the crystalline basement. This makes it difficult to decipher the initial structure of the region, especially if faults of different kinematics and ages are superimposed on each other. In this case, in addition to detailed structural and material studies of fault zones, thermochronological analysis, especially of igneous and metamorphic rocks, which are usually formed in the roots of these orogenies, becomes important.

Thermochronological analysis is based on the use of a set of dating methods for minerals characterized by different closing temperatures [55]. A comparison of the recorded age values of isotopic systems with their closure temperatures, taking into account the range of variations in the temperature gradient of the continental crust (30–12°/km [56]), allows us to obtain an idea of the position of the rock in depth at various time intervals.

A comparison of U/Pb dating of zircon, ⁴⁰Ar/³⁹Ar dating of minerals with different closing temperatures of the K/Ar isotope system (amphibole, mica, feldspar), and Apatite Fission Track (AFT) dating enables an estimate of the thermal history of rocks [57,58], which differs in the case of manifestations of thrust and strike-slip structures. The cover structures caused by thrust processes are characterized by a large vertical component coherent with the shortening and, as a result, a higher cooling rate on the thermal curve reconstructed based on the closure of the sequence of isotope systems and minerals. For strike-slip structures, the predominant component of movement is horizontal, so it is logical to expect that the vertical component of rock movement will not be as high as in the case of thrust structures. In other words, the rock systems of the cover and strike-slip type should be characterized by different denudation rates. Based on this, it can be confidently stated that the thermal history of rocks, for example, granitoid batholiths, containing the largest number of minerals of various closing temperatures of isotope systems, supplemented, if possible, by thermobarometric reconstructions, can serve as an independent source of information about the tectonic evolution of fold regions.

This article is devoted to the reconstruction of the thermal history of the late Paleozoic Kalba–Narym batholith in relation to the manifestation of the Surov gabbroid massif and the Chechek metamorphic complex (granite–gneiss structure), as well as the evolution of the fault tectonics of the ISZ. In order to understand which stages of the ISZ evolution are associated with the formation and further exhumation of the Kalba–Narym batholith, the results of U/Pb isotope dating of zircon, ⁴⁰Ar/³⁹Ar dating of amphibole, biotite, and feldspar, and AFT dating of the listed geological formations are used. According to the authors of the article, the results obtained can make a significant contribution to solving the problem of the tectonic evolution of the AACS and, accordingly, are taken into account in global geodynamic models of the tectonic evolution of the CAFB as a whole.

2. Geological Framework

2.1. The AACS Formation

AACS formed between Siberian and Kazakh paleocontinents [1,2,5,9,10,13] with the closure in the early Carboniferous of the Ob-Zaisan Paleo-Oceanic basin. The final formation of the fold structure occurred at the end of the early Carboniferous (Serpukhov), which is recorded by the appearance of continental molasses deposits of the Bashkir stage with basal conglomerates in localized intermountain depressions. The late Paleozoic active margin of the Siberian continent corresponds to the Rudny–Altai island-arc system, characterized by late Silurian–early Carboniferous volcanogenic sedimentary and granitoid complexes. From the northeast, it is separated by the North-East Fault from the Caledonian Gorny–Altai terrane, and from the southwest, through the ISZ, it is adjacent to the pre-arc turbidite of the Kalba–Narym terrane (Figure 1 and Figure 2). The source of the material for the Kystav–Kurchum (D₂gv) and Takyr (D₃-C₁) formations suites was probably the Rudno–Altaisk island arc system [3,9].

The processes of soft collision interaction of the Siberian and Kazakh continents (not reaching the direct collision of hard continental blocks) caused the deformation of volcanogenic sedimentary rocks of the AACS and their compression, close to the isoclinal folds of the northwestern strike of the axial planes, inclined generally to the northeast [2,10]. In the early Permian, granitoids of the Kalba batholith were formed within the Kalba–Narym terrane and in its marginal part—the ISZ, including the Surov gabbro massif and the Chechek metamorphic complex (Figure 2). The collision of the Kazakh and Siberian plates took place when they rotated clockwise relative to each other, which led to an important structural role of large-amplitude shear displacements at all stages of the AACS development [4,12,13,15,17,18,37,38,46,47,59,60,61,62]. They formed the Zharma, Chara, Irtysh, and Northeastern shear zones (Figure 1). The strike-slip structures were accompanied by less amplitude thrust movements along the feathering faults of the sub-latitudinal orientation.

The largest trans-regional tectonic strike-slip displacement in Central Asia is considered to be the Irtysh Shear Zone (ISZ) [4,12,13,17,18,31,32,33,34,35,36]. Based on structural-petrological and geochronological (⁴⁰Ar/³⁹Ar) studies, two episodes of large-amplitude left-lateral plastic and brittle-plastic deformations with ages of 285–275 and 270–260 Ma have been identified [38]. For the Kalba batholith belt located on the southwestern flank of the ISZ, estimates of the age of formation based on U/Pb dating by zircon fall within the range of 290–274 Ma [38,39]. On this basis, as well as based on structural observations, the sin-tectonic nature of the Kalba batholith belt is assumed [13]. At the same time, the synchronicity of the formation of batholith granitoids with the formation of the Tarim large igneous province [40,41,42,43,44,45] and the close geographical location suggested that the Tarim mantle plume is the most likely energy source that caused the formation of the Kalba–Narym batholith [46,47]. Thus, an interconnected reconstruction of the history of the formation and further tectonic evolution of the ISZ and the Kalba–Narym batholith is relevant.

Throughout the ISZ, deeply metamorphosed tectonic blocks were mapped, for example, Predgornensky, Chechek, and Kurchum (Figure 2a). There are different ideas about the age and genesis of these blocks. Some authors believe that they are exhumed fragments of the Precambrian basement [48,49]. According to another concept, the blocks were formed because of the thermal action of the gabbro of the Irtysh complex, which intruded into the ISZ [50,51]. Within the Chechek structure, the gabbro of the Irtysh complex manifested itself as a large Surov massif (Figure 2).

Figure 2. (a) Scheme of the geological structure of the Irtysh shear zone and the Kalba–Narym batholith according to [39], modified. Bold rectangles with arrows show blocks of deeply metamorphosed rocks within the ISZ: Pr—Predgornensky, Ch—Chechek, and Kr—Kurchum. The numbers in the rectangles are geochronological data (million years old). The dates for zircon (zr) were obtained by U/Pb, for feldspar (fsp), muscovite (mu), and biotite (bt) were obtained by the ⁴⁰Ar/³⁹Ar method, and for apatite by the FT method. The dates obtained in this work are highlighted in green. The letter S in the circle indicates the Sebinsky massif of the Kalba–Narym batholith. (b) Schematic section of the Kalba–Narym terrane, the Irtysh shear zone, and the outskirts of the Rudny Altai terrane.

Figure 2. (a) Scheme of the geological structure of the Irtysh shear zone and the Kalba–Narym batholith according to [39], modified. Bold rectangles with arrows show blocks of deeply metamorphosed rocks within the ISZ: Pr—Predgornensky, Ch—Chechek, and Kr—Kurchum. The numbers in the rectangles are geochronological data (million years old). The dates for zircon (zr) were obtained by U/Pb, for feldspar (fsp), muscovite (mu), and biotite (bt) were obtained by the ⁴⁰Ar/³⁹Ar method, and for apatite by the FT method. The dates obtained in this work are highlighted in green. The letter S in the circle indicates the Sebinsky massif of the Kalba–Narym batholith. (b) Schematic section of the Kalba–Narym terrane, the Irtysh shear zone, and the outskirts of the Rudny Altai terrane.

Although the main stages of the geodynamic evolution of the Kalba–Narym batholith, Rudny–Altai terrane, and the ISZ considered in this article have been generally identified, many unresolved questions remain, for example, about the role of the ISZ in the formation of the Kalba–Narym granitoid batholith, as well as about the role of the ISZ in its subsequent post-magmatic history. There are still different interpretations for the evolution of the kinematics of deformations of the ISZ. In a number of works, the point of view is expressed that the ISZ is an important structure in the framework of the Central Asian Orogenic Belt (CAOB). It represents the site of the final collision of Kazakhstan with Siberia during Hercynian times and records up to 1000 km of lateral displacement during subsequent reorganization in the CAOB edifice [4,12,13,17,18,31].

The authors of the following works believe that the history of ISZ includes three episodes of late Paleozoic deformation representing orogenic thickening, collapse, and transpressional deformation. On a larger scale, late Paleozoic shearing accommodated the eastward migration of internal segments of the western CAOB, possibly associated with the amalgamation of multiple arc systems and continental blocks [22,25,33,35,36]. The systematic reconstruction of the cooling history of the granitoid massifs of the Kalba–Narym batholith is still lacking, as well as a model for its denudation.

2.2. The Kalba–Narym Batholith

The granitoids of the Kalba–Narym zone, forming one of the largest batholiths in the western Central Asian fold belt, occupy a significant part of the area of the Kalba–Narym terrane (>40,000 km²). According to geophysical data, the thickness of granitoid massifs ranges from 2 to 12 km with a predominance of 7–10 km [63,64].

The granitoid Kalba–Narym batholith includes the Kalgutin granodiorite–granite complex, and the Kunush–plagiogranite complex, the Kalba granodiorite–granite complex, the Monastery granite–leucogranite complex. The batholith is intruded by “post-batholith” dikes of the Mirolyubov complex. Also within the Kalba–Narym terrane near Ust-Kamenogorsk are gabbro of the Surov massif of the Irtysh complex and metamorphic rocks of the Chechek granite–gneiss structure, distinguished as part of the ISZ (Figure 2) [46,60,61,65,66,67].

2.3. The Irtysh Gabbro Complex

In the marginal part of the Kalba–Narym terrane, within the ISZ, outcrops of gabbro of the Surov massif (313 ± 1 Ma) [61]) close in age with the Chechek complex are found (312 ± 3 Ma) [12,65,66]. The origin of metamorphic rocks is directly related to the episode of the introduction and formation of the Surov gabbro lopolith, which provided the necessary heating and melting of the overlying strata, and after consolidation (312 ± 3 Ma ago)—reservation and “protection” from late (~280 and ~260 Ma) large-scale shear deformations along the ISZ [65,66].

Rocks of the early Carboniferous Irtysh gabbro complex, of which the Surov massif is a part [61,68,69], intrude middle Devonian sedimentary rocks (Kystav–Kurchum and Takyr suites). Their contacts are intrusive in the case of large gabbro bodies and predominantly tectonic in the case of small bodies. The small bodies of gabbro are most often rootless tectonic sheets or boudin-like bodies among metamorphic rocks. The Surov intrusion is lopolite shaped (Figure 2 and Figure 3) and has mainly NE dipping magmatic layering. Its basal series exposed in the southern part of the intrusion consists of pre dominant melanocratic gabbro and minor amounts of ultramafic rocks (mainly olivine cumulate).

2.4. The Chechek Metamorphic Complex

Near Ust’-Kamenogorsk, the Surov gabbro is adjacent to the gneiss and diatectite of the Chechek metamorphic complex of epidote–amphibolite to amphibolite facies [65]. The core of this complex consists mainly of granite gneiss and numerous synmetamorphic granite veins. The presence of garnet, sillimanite, and cordierite in granite gneiss, as well as granoblastic and coronite textures, records a prolonged heating effect. Minerals in the center of the Chechek complex formed at T = 665–720 °C and P = 0.4–0.6 GPa, in the upper amphibolite facies conditions. Thus, the metamorphism was a contact one and was a consequence of the thermal effect of gabbro [65].

3. Materials and Methods

Geological field research was conducted mostly through hiking trails, documenting rock relationships, sketching key outcrops, and sampling. The testing was carried out on the bedrock outcrops or on the blocky ruins of the same rocks directly on the bedrock outcrops. The freshness of the selected samples was visually controlled by the gloss of micas, amphiboles, etc. Hand-held GPS receivers were used for coordinate control of the sampling points. Samples were taken from various parts of the cover-thrust structure of the region. The most important of them are the granitoid complexes of the Kalba–Narym terrane, the Surov gabbro massif, and the Chechek metamorphic complex. Whenever possible, rock samples were taken that contain the maximum amount of various minerals (zircon, apatite, amphibole, mica, feldspars), differing in the closing temperatures of isotopic systems. The obtained results of the dating of these minerals eventually made it possible to create a tectonothermal reconstruction of the cover-thrust structure.

In order to reconstruct the history of the formation of the cover-thrust structure of the Kalba–Narym terrane and, in general, the AACS, we compared thermochronological data for granitoid complexes of the Kalba–Narym batholith, as well as for the Surov Gabbro massif and the Chechek metamorphic complex. Eleven samples of granitoid complexes of the Kalba–Narym batholith were used, which were previously characterized in detail using petrological and geochemical data, and for which the age of zircon was determined by the U/Pb method [38]. Monofractions of biotite and muscovite were selected from a sample of cataclased gabbro (sample B-23-146).

The thermochronological approach is based on the kinetic parameters of daughter isotopes measured in laboratory conditions and assumes the immutability of the crystal structure of the mineral while considering the case when the closure temperature is significantly lower than the temperature of the formation of the mineral—T_c < T_f. In this case, the obtained dating corresponds to the closure time of the corresponding isotope system. In some situations, the ratio between the closure temperature of the isotope system T_c and the formation temperature T_f of the dated mineral phase is the opposite—T_c > T_f. For example, the U/Pb dates for zircon correspond to this case. With this ratio, the obtained dating directly corresponds to the age of formation of the dated mineral phase, namely, zircon.

⁴⁰Ar/³⁹Ar biotite dating was performed on 11 samples (B-23-146, K-14-16, X-1041, X-1042, X-1044, X-1045, X-1047, X-1052, X-1056, 2463, 2458), whereas feldspar was used only once (sample K-14-19). Sample B-23-146 was instead dated with both biotite and muscovite. Petrological and petrochemical descriptions of the samples are given in the relevant publications (see

Table 1. Summary of thermochronological data for granitoids of the Kalba–Narym batholith and the Chechek granite–gneiss structure.

Sample/ Massif	Complex/ Rock/Mineral *	Method **	Age (Ma)	Closure/ Formation T (°C) ***	Reference
X-1052	Kalgutin/ granodiorite/zrn	U/PbL	308 ± 2	~850f	[38]
X-1052 Kurchum	Kalgutin/ granodiorite/bt	40Ar/39Ar	289 ± 3	330c	This work
X-1047	Kalgutin/ granodiorite/zrn	U/PbL	303 ± 1	~850f	[38]
X-1047 Kurchum	Kalgutin/ granodiorite/bt	40Ar/39Ar	282 ± 3	330c	This work
X-1056	Kalba/granite/zrn	U/PbL	297 ± 1	~806f	[38]
X-1056/ Asubulak	Kalba/granite/bt	40Ar/39Ar	285 ± 2	330c	This work
X-1045	Kalba/granodiorite/zrn	U/PbL	297 ± 1	~806f	[38]
X-1045 Chernovin		40Ar/39Ar	281 ± 2	330c	This work
X-1042	Kalba/granite/zrn	U/PbL	286 ± 3	~806f	[38]
X-1042 Chernovin	Kalba/granite/bt	40Ar/39Ar	268 ± 2	330c	This work
Narym	Kalba/granite/zrn	U/PbL	296 ± 4	~806f	[5]
2458	Kalba/granite/bt	40Ar/39Ar	275 ± 3	330c	This work
2463	Kalba/granite/bt	40Ar/39Ar	283 ± 3	330c	This work
X-1044	Kalba/granite/zrn	U/PbL	288 ± 1	~806f	[38]
X-1044 Chernovin	Kalba/granite/bt	40Ar/39Ar	282 ± 2	330c	This work
X-1041	Monastery/ leucogranite/zrn	U/PbL	283 ± 2	~810f	[38]
X-1041 Voylochev	Monastery/ leucogranite/bt	40Ar/39Ar	285 ± 2	330c	This work
8-03-10 Sebin	Monastery/ leucogranite/zrn	U/PbL	284 ± 4	~810f	[38]
KA-14-18	Monastery/ leucogranite/bt	40Ar/39Ar	280 ± 2	330c	This work
KA-14-18 Sebin	Monastery/ leucogranite/fsp	40Ar/39Ar	243 ± 3	250c	This work
KA-21-345 KA-21-346 KA-21-347 Sebin	Monastery/ leucogranite/ap	FST	Average 229 ± 21	110c	[70]
X-1414 Surov intrusion	Irtysh/gabbronorite/zrn	U/PbL	313 ± 1	>900f	[61]
X-1207 Surov intrusion	Irtysh/gabbrodiorite/zrn	U/PbL	313 ± 1	>900f	[61]
E-32 Chechek	Chechek/granite-gneiss/ms	40Ar/39Ar	312 ± 3	360c	[65]
KZ-06 Chechek	Chechek/granite-gneiss/ap	FST	87 ± 6	110c	[17]
B-23-146	Chechek/gabbro tectonite/ms	40Ar/39Ar	304 ± 4	360c	This work
B-23-146 Surov intrusion	Chechek/gabbro tectonite/bt	40Ar/39Ar	289 ± 3	330c	This work

* The following designations are used in the table: zrn—zircon, ms—muscovite, bt—biotite, fsp—feldspar, ap—apatite. ** U/Pb zircon dating was performed using the ICP laser ablation mass spectrometry method (U/Pb_L). *** T_c is the closing temperature of the corresponding isotope system; T_f is the temperature of formation of the corresponding mineral phase.

). Minerals for ⁴⁰Ar/³⁹Ar dating were extracted by the conventional techniques of magnetic and density separation at the V.S. Sobolev Institute of Geology and Mineralogy (Novosibirsk, Russia).

⁴⁰Ar/³⁹Ar dating of monomineral fractions was carried out at the Analytical Center for Multi-element and Isotope Studies (Novosibirsk, Russia) as in [57]. Samples were wrapped in aluminum foil together with biotite MCA-11 standard monitor samples, placed in quartz capsules, vacuumed, and sealed. For muscovite MCA-11, as a result of its calibration using muscovite Bern 4m and biotite LP-6, we assumed an average age value of 311.0 ± 1.5 Ma. The error of the calculated value of J, characterizing the neutron flux value for each of the positions within the ampoule, did not exceed 0.5%. The capsules were irradiated by fast neutrons in a Cd-lined tube of the IRT-T nuclear reactor at Tomsk Polytechnical University, in Tomsk, Russia. The samples were subjected to stepwise heating in a quartz tube using an external tubular resistance furnace. A chromel-alumel thermocouple was used to control the temperature with the accuracy of ±3 °C. The ⁴⁰Ar and ³⁶Ar blank runs (10 min at 1200 °C) did not exceed 3 × 10⁻¹⁰ and 0.003 × 10⁻¹⁰ ncm³, respectively. Argon cleaning was performed using ZrAl-SAES getters (SAES Advanced Technologies Spa, Avezzano, Italy). The argon isotope composition was measured on a Micromass Noble Gas 5400 mass spectrometer (Micromass UK Limited, Wilmslow, UK), to an accuracy of ±1σ. The contribution of interfering Ar isotopes formed together with ³⁹Ca and ⁴⁰K was estimated using the coefficients of (³⁹Ar/³⁷Ar)_Ca = 0.001279 ± 0.000061, (³⁶Ar/³⁷Ar)_Ca = 0.000613 ± 0.000084, and (⁴⁰Ar/³⁹Ar)_K = 0.0191 ± 0.0018. The plateau ages were calculated in Isoplot-3 [71], as weighted average values over at least three successive temperature steps. The results were interpreted with reference to the conventional criteria [72]. Analytical error is specified at the level of 1σ.

4. Results

4.1. The Fieldwork

The fieldwork of our research over the past few years revealed that in the area of Ust-Kamenogorsk, the boundary of the Irtysh gabbro complex and the metamorphic rocks of the Chechek complex is represented by a thrust (Figure 3). The plunging of the thrust planes is defined as north-eastern. They are transformed into dome-shaped folds. The thrust structure is everywhere disrupted by right-lateral shifts of north-eastern strike, deforming it into Z-shaped folds with amplitudes of up to a few tens of cm. The shear deformations are most fully manifested at the boundary with the Rudny–Altai terrane, where they are represented by a zone of green schists with a thickness of up to many hundreds of meters.

Figure 3. The scheme of the geological structure of the Irtysh shear zone in the area of Ust-Kamenogorsk.

Figure 3. The scheme of the geological structure of the Irtysh shear zone in the area of Ust-Kamenogorsk.

At the base of the thrust (Figure 3), there is a zone in which biotite–muscovite shales include lenticular boudins of cataclastyc gabbro. They are jointly crushed into isoclinal folds. Structurally, blastomylonites, crystalline schists, and gneisses of metapelite composition (biotite, garnet–biotite, biotite–garnet–sillimanite) and granite–gneisses of the Chechek metamorphic complex are located below the Surov massif. The complex is deformed into dome-shaped folds and pushed over sedimentary rocks of the Takyr formation. In the upper part of the metamorphic complex, there are spherical gabbro blocks with a diameter of several meters, in the lower part, there are tectonic lenses of metasedimentary rocks, Kystav–Kurchum Fms.

The thickness of the tectonic lenses reaches 1–2 m, the length is several meters (Figure 4). Figure 4a shows the general appearance of the tectonic mélange. The matrix is represented by granite gneisses, which include oval blocks of quartz–biotite, garnet–biotite, and quartz–biotite–garnet–sillimanite schists of various sizes (up to 70 cm).

In Figure 4b, among the granite gneisses, there is a block of rhythmically layered rocks, the width and height of which are about 80 cm. The layering in the block is oriented at a right angle to the strike of the metamorphic schistosity in the granite gneisses. It is expressed by the alternation of quartz–biotite (quartz sandstones), biotite, and siliceous layers, which form a rhythmic alternation. The thickness of grey and light green siliceous rocks and grey quartz–biotite interlayers is from a few mm to several cm. Figure 4c shows isoclinal folding in granite–gneisses, including blocks of quartz–biotite and siliceous rocks. Black biotite zones are clearly visible in granite–gneisses. The amplitude of isoclinal folding is about 30 cm, the subsidence of axial planes is consistent with the general sink of granite–gneisses to the northeast, the subsidence angle is 45°. Figure 4d shows a large block of metasedimentary rocks 3 m long and 1 m high. It is represented by alternating quartz–biotite, biotite, and siliceous interlayers. The rocks are intersected by a quartz vein, which creates the appearance of a recumbent fold. Metasedimentary rocks subside to the north-northeast at an angle of 40°.

The metamorphic rocks of the Chechek complex, taking into account the superimposed dome-shaped folding, are characterized by a gentle northeastern immersion of the axial planes of the drawing folds and mineral linearity in metamorphic rocks. Its thickness is estimated at 1200 m. In general, the metamorphic rocks of the Chechek complex and Surov gabbro should be considered as a tectonic mélange formed by the framing of a large gabbro body intruded in the thrust. The southern cataclased gabbro contact with metamorphic rocks sinks to the north at angles of 40–50°, and the eastern and western contacts are gently sloping to the northeast-east. In places, hot contacts of gabbro and metamorphic rocks have been preserved [61,65,68,69].

Tectonically, above the massive gabbro on the right bank of the Irtysh River, there are also metamorphic rocks (mainly blastomylonites in crystalline shales and granite–gneisses), including bodies of gabbro cataclasite up to the first meters in size. Even higher are the metamorphic rocks of the green schist facies of metamorphism, formed by middle Devonian sedimentary rocks (Kystav–Kurchum Fms.). They are characterized by frequent interlayers of siliceous rocks and the mineral linearity of muscovite, manifested along the planes of axial cleavage. The layers are often crumpled into shallow recumbent folds with axial planes and mineral linearity sinking to the northeast and east with angles of 35–10°.

First of all, they often manifest themselves in Z-shaped folds with vertically plunging hinges. The folds are formed in narrow fault zones represented by green schist, mylonites, and blastomylonites from the rocks of the Irtysh metamorphic complex of the epidote amphibolite facies of metamorphism.

The metamorphic shear rocks are most fully represented along the northeastern border of the ISZ on the border with the Rudny Altai terrane (Figure 2 and Figure 3). In the frontal part of the thrust, the granitoids of the Kalba–Narym batholith are located, embedded in the middle Devonian sedimentary rocks (Kystav–Kurchum Fms.) (Figure 2).

4.2. The ⁴⁰Ar/³⁹Ar Dating

In the ⁴⁰Ar/³⁹Ar spectra of almost all analyzed mineral fractions (

Table A1. ⁴⁰Ar/³⁹Ar data for minerals from the samples of the Kalba–Narym batholith and the Chechek granite–gneiss structure.

T °C	40Ar (cm3 STP)	40Ar/39Ar ± 1σ	38Ar/39Ar ± 1σ	37Ar/39Ar ± 1σ	36Ar/39Ar ± 1σ	Ca/K	∑39Ar (%)	Age (Ma) ± 1σ	40Ar * (%)
Sample X-1041 biotite, weight 4.95 mg, J = 0.004528 ± 0.000054 1, plateau age (850–1100 °C) = 284.8 ± 3.8 Ma (1σ)
500	390.5	57.3 ± 0.3	0.01922 ± 0.00520	0.00002 ± 0.00002	0.00021 ± 0.00436	0.0001	4.6	291.7 ± 6.6	99.9
650	2617.0	53.4 ± 0.1	0.00770 ± 0.00035	0.00008 ± 0.00003	0.00001 ± 0.00083	0.0003	37.6	273.5 ± 2.4	100.0
750	1191.1	54.2 ± 0.1	0.01199 ± 0.00168	0.04209 ± 0.00460	0.00001 ± 0.00169	0.1515	52.4	277.6 ± 3.2	100.0
850	2415.3	57.0 ± 0.1	0.00001 ± 0.00035	0.00026 ± 0.00250	0.00001 ± 0.00069	0.0009	81.0	290.6 ± 2.4	100.0
950	853.8	55.7 ± 0.2	0.00523 ± 0.00211	0.00009 ± 0.00007	0.00014 ± 0.00272	0.0003	91.3	284.2 ± 4.4	99.9
1025	234.8	54.0 ± 0.3	0.00049 ± 0.00007	0.00027 ± 0.00031	0.00044 ± 0.00554	0.0010	94.3	275.8 ± 8.2	99.8
1100	461.2	54.3 ± 0.1	0.00826 ± 0.00263	0.00107 ± 0.00002	0.00033 ± 0.00149	0.0039	100.0	277.4 ± 3.0	99.8
Sample KA-14-16 biotite, weight 52.01 mg, J = 0.004528 ± 0.000054 1, plateau age (720–1130 °C) = 279.1 ± 3.3 Ma (1σ)
500	22.5	43.8 ± 0.7	0.05938 ± 0.02329	16.63285 ± 7.78847	0.06168 ± 0.01523	59.9	0.7	162.8 ± 27.5	58.4
610	97.2	59.8 ± 0.3	0.05926 ± 0.00630	3.99485 ± 2.93448	0.08588 ± 0.00393	14.4	3.0	216.0 ± 7.2	57.6
720	371.4	48.0 ± 0.1	0.01561 ± 0.00161	1.73254 ± 0.60034	0.00884 ± 0.00130	6.3	13.6	279.6 ± 3.4	94.6
800	844.9	46.4 ± 0.1	0.01708 ± 0.00025	0.01537 ± 0.15000	0.00280 ± 0.00051	0.1	38.8	280.9 ± 2.7	98.2
850	509.1	46.7 ± 0.1	0.01585 ± 0.00100	0.14254 ± 0.38051	0.00485 ± 0.00090	0.5	53.8	279.1 ± 3.0	96.9
920	296.8	47.4 ± 0.1	0.02001 ± 0.00191	1.82078 ± 1.74010	0.00974 ± 0.00179	6.5	62.4	274.9 ± 3.9	93.9
1020	730.9	46.6 ± 0.1	0.01493 ± 0.00108	0.02376 ± 0.30843	0.00487 ± 0.00091	0.1	84.1	278.5 ± 3.0	96.9
1080	416.8	47.3 ± 0.1	0.01827 ± 0.00208	2.03645 ± 0.75456	0.00751 ± 0.00130	7.3	96.2	278.2 ± 3.4	95.3
1130	131.2	48.0 ± 0.1	0.01356 ± 0.00322	3.15342 ± 1.56345	0.00651 ± 0.00311	11.3	100.0	283.6 ± 5.9	96.0
Sample KA-14-19 feldspar, weight 125.67 mg, J = 0.004528 ± 0.000054 1, plateau age (825–1130 °C) = 245.9 ± 3.1 Ma (1σ)
500	16.9	65.7 ± 1.7	0.02802 ± 0.02545	1.26424 ± 3.96484	0.05341 ± 0.01919	4.5	0.4	301.8 ± 32.4	76.0
625	240.2	40.0 ± 0.1	0.01436 ± 0.00227	0.56055 ± 0.24672	0.01465 ± 0.00147	2.0	9.8	220.7 ± 3.2	89.2
725	524.4	40.4 ± 0.1	0.01805 ± 0.00072	0.40446 ± 0.24077	0.00672 ± 0.00034	1.5	30.1	236.5 ± 2.2	95.1
825	517.4	41.2 ± 0.1	0.01845 ± 0.00062	0.35746 ± 0.18756	0.00590 ± 0.00109	1.3	49.8	243.0 ± 2.9	95.8
925	515.0	41.7 ± 0.1	0.01629 ± 0.00125	0.31492 ± 0.28741	0.00517 ± 0.00078	1.1	69.1	246.9 ± 2.6	96.3
1025	429.3	42.0 ± 0.1	0.01733 ± 0.00101	0.69472 ± 0.04437	0.00849 ± 0.00136	2.5	85.1	243.2 ± 3.2	94.0
1130	405.6	42.7 ± 0.1	0.01868 ± 0.00158	1.80304 ± 0.54957	0.00575 ± 0.00157	6.5	100.0	251.4 ± 3.5	96.0
Sample X-1045 biotite, weight 9.86 mg, J = 0.004528 ± 0.000054 1, plateau age (750–1100 °C) = 260.4 ± 3.0 Ma (1σ)
500	91.0	45.9 ± 0.3	0.06044 ± 0.01423	3.80707 ± 0.18499	0.02491 ± 0.00748	13.70	0.4	198.8 ± 11.0	84.0
600	965.2	55.2 ± 0.1	0.02026 ± 0.00145	0.25325 ± 0.02180	0.00471 ± 0.00109	0.91	3.7	285.2 ± 2.6	97.5
700	6005.1	53.1 ± 0.1	0.01943 ± 0.00024	0.03393 ± 0.00338	0.00196 ± 0.00053	0.12	25.4	266.0 ± 2.1	98.9
750	5037.7	51.8 ± 0.1	0.01511 ± 0.00008	0.01110 ± 0.00074	0.00174 ± 0.00058	0.04	44.1	260.3 ± 2.1	99.0
800	1765.3	51.7 ± 0.1	0.01582 ± 0.00033	0.00828 ± 0.00220	0.00083 ± 0.00035	0.03	50.6	260.9 ± 2.0	99.5
900	1522.2	51.4 ± 0.1	0.01543 ± 0.00024	0.02381 ± 0.00194	0.00268 ± 0.00046	0.09	56.3	257.1 ± 2.0	98.5
1025	6418.3	51.9 ± 0.1	0.01449 ± 0.00008	0.00698 ± 0.00046	0.00123 ± 0.00054	0.02	80.1	261.2 ± 2.1	99.3
1100	5418.4	52.1 ± 0.1	0.01400 ± 0.00009	0.00207 ± 0.00067	0.00106 ± 0.00031	0.01	100.0	262.6 ± 2.0	99.4
Sample X-1056 biotite, weight 71.5 mg, J = 0.004528 ± 0.000054 1, plateau age (650–1130 °C) = 284.6 ± 3.3 Ma (1σ)
500	71.2	42.3 ± 0.1	0.02143 ± 0.00369	0.14170 ± 0.01263	0.03549 ± 0.00165	0.51	1.3	169.0 ± 2.8	75.2
600	531.1	58.4 ± 0.1	0.01791 ± 0.00040	0.02686 ± 0.00165	0.01668 ± 0.00036	0.10	8.4	275.7 ± 2.1	91.6
650	961.7	56.8 ± 0.1	0.01561 ± 0.00020	0.00895 ± 0.00096	0.00504 ± 0.00026	0.03	21.6	284.7 ± 2.2	97.4
700	1151.4	56.1 ± 0.1	0.01555 ± 0.00018	0.00523 ± 0.00080	0.00246 ± 0.00023	0.02	37.6	285.1 ± 2.2	98.7
750	623.5	55.7 ± 0.1	0.01624 ± 0.00013	0.01362 ± 0.00144	0.00220 ± 0.00056	0.05	46.3	283.5 ± 2.3	98.8
850	229.2	55.9 ± 0.1	0.01740 ± 0.00110	0.02538 ± 0.00314	0.00093 ± 0.00096	0.09	49.5	286.1 ± 2.5	99.5
975	2125.4	56.2 ± 0.1	0.01544 ± 0.00008	0.01527 ± 0.00029	0.00235 ± 0.00025	0.05	78.9	285.7 ± 2.2	98.8
1050	1223.3	55.9 ± 0.1	0.01539 ± 0.00019	0.03660 ± 0.00062	0.00218 ± 0.00032	0.13	96.0	284.4 ± 2.2	98.9
1130	288.9	56.0 ± 0.1	0.01608 ± 0.00041	0.15098 ± 0.00084	0.00354 ± 0.00061	0.54	100.0	282.7 ± 2.3	98.1
Sample X-1044 biotite, weight 71.5 mg, J = 0.004528 ± 0.000054 1, plateau age (650–1130 °C) = 284.6 ± 3.3 Ma (1σ)
500	19.0	39.2 ± 0.2	0.02941 ± 0.00265	0.40784 ± 0.01722	0.03673 ± 0.00212	1.47	0.6	150.7 ± 3.5	72.3
600	140.4	49.1 ± 0.1	0.01934 ± 0.00098	0.09718 ± 0.00144	0.01826 ± 0.00049	0.35	4.4	227.9 ± 1.9	89.0
700	934.6	55.3 ± 0.1	0.01565 ± 0.00017	0.02390 ± 0.00042	0.00370 ± 0.00018	0.09	26.9	278.5 ± 2.1	98.0
775	599.7	55.7 ± 0.1	0.01575 ± 0.00028	0.01383 ± 0.00187	0.00173 ± 0.00053	0.05	41.2	282.9 ± 2.3	99.1
875	658.0	56.0 ± 0.1	0.01646 ± 0.00041	0.03710 ± 0.00134	0.00374 ± 0.00039	0.13	56.9	281.6 ± 2.2	98.0
975	554.3	56.2 ± 0.1	0.01502 ± 0.00032	0.04986 ± 0.00169	0.00263 ± 0.00030	0.18	70.0	284.0 ± 2.2	98.6
1050	749.2	55.9 ± 0.1	0.01536 ± 0.00033	0.07133 ± 0.00117	0.00257 ± 0.00016	0.26	87.8	282.8 ± 2.1	98.6
1130	512.0	56.0 ± 0.1	0.01504 ± 0.00023	0.19017 ± 0.00162	0.00416 ± 0.00035	0.68	100.0	281.0 ± 2.2	97.8
Sample X-1042 biotite, weight 5.81 mg, J = 0.004528 ± 0.000054 1, plateau age (900–1100 °C) = 268.3 ± 2.1 Ma (1σ)
500	500.0	40.8 ± 0.1	0.02330 ± 0.00072	0.03415 ± 0.00579	0.02496 ± 0.00101	0.12	3.8	175.4 ± 2.0	81.9
600	3636.3	51.4 ± 0.1	0.01567 ± 0.00013	0.00354 ± 0.00103	0.00405 ± 0.00027	0.01	25.5	257.6 ± 2.0	97.7
675	2827.7	52.2 ± 0.1	0.01543 ± 0.00016	0.00801 ± 0.00123	0.00160 ± 0.00049	0.03	42.1	264.6 ± 2.1	99.1
800	2231.2	52.3 ± 0.1	0.01390 ± 0.00019	0.01403 ± 0.00131	0.00352 ± 0.00030	0.05	55.2	262.6 ± 2.0	98.0
900	1910.8	53.0 ± 0.1	0.01665 ± 0.00019	0.01726 ± 0.00161	0.00566 ± 0.00048	0.06	66.3	262.6 ± 2.1	96.8
1000	3810.1	53.0 ± 0.1	0.01522 ± 0.00012	0.01699 ± 0.00119	0.00161 ± 0.00053	0.06	88.3	268.5 ± 2.2	99.1
1050	1336.5	53.3 ± 0.1	0.01608 ± 0.00029	0.04024 ± 0.00315	0.00296 ± 0.00036	0.14	96.0	268.2 ± 2.1	98.4
1100	698.2	53.8 ± 0.1	0.02001 ± 0.00074	0.20121 ± 0.00596	0.00454 ± 0.00075	0.72	100.0	268.4 ± 2.3	97.5
Sample 2463 biotite, weight 21.75mg, J = 0.004528 ± 0.000054 1, plateau age (900–1130 °C) = 283.3 ± 2.9 Ma (1σ)
500	30.0	33.3 ± 0.2	0.02120 ± 0.02120	0.08281 ± 0.01910	0.03934 ± 0.00628	0.2981	1.4	157.4 ± 13.0	65.2
600	124.5	35.0 ± 0.1	0.02333 ± 0.02333	0.02432 ± 0.00227	0.02595 ± 0.00105	0.0876	7.0	196.5 ± 2.9	78.2
700	603.9	42.3 ± 0.1	0.02018 ± 0.02018	0.01386 ± 0.00076	0.01009 ± 0.00041	0.0499	29.3	275.7 ± 2.9	93.0
800	578.4	43.5 ± 0.1	0.02021 ± 0.02021	0.01265 ± 0.00103	0.00903 ± 0.00053	0.0455	50.0	285.7 ± 3.1	93.9
900	278.7	43.2 ± 0.1	0.02146 ± 0.02146	0.02131 ± 0.00300	0.00928 ± 0.00077	0.0767	60.1	283.5 ± 3.2	93.7
1000	542.6	43.3 ± 0.1	0.02125 ± 0.02125	0.02368 ± 0.00146	0.01160 ± 0.00082	0.0852	79.6	279.2 ± 3.3	92.1
1130	564.7	43.3 ± 0.1	0.01970 ± 0.01970	0.07923 ± 0.00085	0.00912 ± 0.00028	0.2852	100.0	284.2 ± 2.9	93.8
Sample 2458 biotite, weight 39.8 mg, J = 0.004528 ± 0.000054 1, plateau age (800–1130 °C) = 275.2 ± 3.4 Ma (1σ)
500	72.6	33.2 ± 0.1	0.02810 ± 0.00232	0.06164 ± 0.00328	0.03617 ± 0.00272	0.22	1.7	159.5 ± 5.7	67.8
600	257.9	37.1 ± 0.1	0.02130 ± 0.00044	0.02316 ± 0.00169	0.02176 ± 0.00069	0.08	7.0	214.4 ± 2.6	82.7
700	1329.7	42.3 ± 0.1	0.01945 ± 0.00012	0.01324 ± 0.00040	0.01173 ± 0.00020	0.05	31.0	267.5 ± 2.7	91.8
800	939.7	43.7 ± 0.1	0.01900 ± 0.00022	0.01347 ± 0.00059	0.01176 ± 0.00020	0.05	47.5	276.0 ± 2.8	92.1
900	863.0	43.1 ± 0.1	0.01906 ± 0.00018	0.01866 ± 0.00060	0.01204 ± 0.00031	0.07	62.8	272.1 ± 2.8	91.8
1000	1175.2	43.5 ± 0.1	0.01941 ± 0.00014	0.01991 ± 0.00042	0.01220 ± 0.00009	0.07	83.4	274.2 ± 2.7	91.7
1130	954.1	44.1 ± 0.1	0.01902 ± 0.00023	0.11439 ± 0.00069	0.01178 ± 0.00026	0.41	100.0	278.5 ± 2.8	92.1
Sample X-1052 biotite, weight 3.3 mg, J = 0.004528 ± 0.000054 1, plateau age (850–1100 °C) = 289.4 ± 3.4 Ma (1σ)
500	20.9	62.2 ± 1.3	0.05037 ± 0.02094	0.87014 ± 0.18577	0.19526 ± 0.02100	3.132	0.6	24.4 ± 32.6	7.4
600	57.1	50.7 ± 0.2	0.03431 ± 0.00644	0.14209 ± 0.06394	0.11289 ± 0.00382	0.511	2.4	91.7 ± 5.8	34.3
700	294.1	58.4 ± 0.1	0.02949 ± 0.00139	0.05529 ± 0.01259	0.01127 ± 0.00136	0.199	10.8	276.5 ± 2.8	94.3
850	1430.0	60.3 ± 0.1	0.01715 ± 0.00023	0.02236 ± 0.00224	0.00558 ± 0.00032	0.080	50.3	293.0 ± 2.2	97.3
950	469.1	60.4 ± 0.1	0.01869 ± 0.00086	0.10875 ± 0.00882	0.01643 ± 0.00098	0.391	63.2	278.5 ± 2.5	92.0
950	469.1	60.4 ± 0.1	0.01869 ± 0.00086	0.10875 ± 0.00882	0.01643 ± 0.00098	0.391	63.2	278.5 ± 2.5	92.0
1025	614.1	60.2 ± 0.1	0.02091 ± 0.00062	0.00100 ± 0.00719	0.00849 ± 0.00087	0.004	80.2	288.5 ± 2.4	95.8
1100	719.8	60.4 ± 0.1	0.01887 ± 0.00070	0.02122 ± 0.00600	0.00381 ± 0.00065	0.076	100.0	296.0 ± 2.4	98.1
Sample X-1047 biotite, weight 5.24 mg, J = 0.004528 ± 0.000054 1, plateau age (800–1100 °C) = 282.4 ± 3.0 Ma (1σ)
500	61.3	61.8 ± 0.5	0.03026 ± 0.00880	0.20401 ± 0.03479	0.00451 ± 0.01198	0.734	0.6	302.0 ± 16.6	97.8
650	1270.7	55.0 ± 0.1	0.00836 ± 0.00019	0.00036 ± 0.00135	0.00853 ± 0.00066	0.001	14.9	265.1 ± 2.2	95.4
700	1563.9	53.6 ± 0.1	0.01071 ± 0.00037	0.01784 ± 0.00141	0.00561 ± 0.00054	0.064	32.9	262.5 ± 2.1	96.9
750	1522.4	53.7 ± 0.1	0.01059 ± 0.00047	0.01534 ± 0.00308	0.00284 ± 0.00074	0.055	50.4	266.8 ± 2.2	98.4
800	501.3	63.8 ± 0.4	0.09076 ± 0.01452	0.00591 ± 0.02531	0.01590 ± 0.01223	0.021	55.2	295.9 ± 16.9	92.7
925	850.6	58.8 ± 0.1	0.00911 ± 0.00225	0.00423 ± 0.00551	0.01633 ± 0.00282	0.015	64.1	272.2 ± 4.4	91.8
1025	1485.6	58.0 ± 0.2	0.00832 ± 0.00177	0.00721 ± 0.00663	0.00321 ± 0.00213	0.026	79.9	286.1 ± 3.7	98.4
1100	1942.4	59.8 ± 0.1	0.01366 ± 0.00172	0.00397 ± 0.00436	0.01070 ± 0.00154	0.014	100.0	284.2 ± 3.0	94.7
Sample B-23-146 biotite, weight 56.52 mg, J = 0.004528 ± 0.000054 1, plateau age (540–1130 °C) = 289.1 ± 3.5 Ma (1σ)
500	1475.3	45.5 ± 0.1	0.05481 ± 0.00154	0.08531 ± 0.00183	0.02000 ± 0.00236	0.307	14.4	272.2 ± 5.3	87.0
540	739.0	46.4 ± 0.1	0.05862 ± 0.00128	0.08216 ± 0.00440	0.01834 ± 0.00234	0.296	21.5	281.2 ± 5.2	88.3
580	2019.2	47.7 ± 0.1	0.03036 ± 0.00164	0.05464 ± 0.00293	0.01642 ± 0.00079	0.197	40.3	292.9 ± 3.3	89.8
680	1897.9	49.0 ± 0.1	0.03041 ± 0.00095	0.04899 ± 0.00194	0.02329 ± 0.00145	0.176	57.6	288.4 ± 4.0	86.0
750	329.6	51.9 ± 0.2	0.06585 ± 0.00199	0.14083 ± 0.00598	0.03874 ± 0.00306	0.507	60.4	278.3 ± 6.4	78.0
850	726.2	48.7 ± 0.2	0.05198 ± 0.00154	0.08909 ± 0.00517	0.02158 ± 0.00330	0.322	67.0	289.7 ± 6.9	86.9
950	898.5	48.3 ± 0.1	0.04831 ± 0.00068	0.07979 ± 0.00307	0.02287 ± 0.00132	0.287	75.3	285.2 ± 3.8	86.0
1050	1654.5	47.6 ± 0.1	0.02852 ± 0.00071	0.04228 ± 0.00164	0.01602 ± 0.00143	0.152	90.7	293.2 ± 4.0	90.1
1090	804.4	47.7 ± 0.1	0.04002 ± 0.00127	0.06076 ± 0.00257	0.01650 ± 0.00153	0.219	98.2	293.1 ± 4.1	89.8
1130	194.5	48.5 ± 0.4	0.13083 ± 0.00451	0.16040 ± 0.01004	0.02136 ± 0.00868	0.577	100.0	288.8 ± 16.6	87.0
Sample B-23-146 muscovite, weight 56.52 mg, J = 0.004528 ± 0.000054 1, plateau age (540–1130 °C) = 289.1 ± 3.5 Ma (1σ)
550	58.1	70.4 ± 1.0	0.07549 ± 0.01069	5.97844 ± 0.15611	0.19980 ± 0.01286	21.522	1.0	81.0 ± 26.2	16.3
650	152.0	106.8 ± 0.5	0.02627 ± 0.00522	0.67446 ± 0.03905	0.23315 ± 0.00191	2.428	2.8	256.8 ± 5.0	35.6
750	478.5	66.7 ± 0.1	0.01776 ± 0.00310	0.06931 ± 0.03274	0.07116 ± 0.00179	0.249	11.9	304.7 ± 4.4	68.5
825	906.9	60.9 ± 0.1	0.01377 ± 0.00175	0.05460 ± 0.01719	0.05143 ± 0.00134	0.197	30.8	305.3 ± 3.9	75.1
910	745.1	55.6 ± 0.2	0.01799 ± 0.00339	0.00549 ± 0.01820	0.03006 ± 0.00323	0.020	47.8	311.3 ± 6.7	84.0
985	1011.3	55.5 ± 0.1	0.01475 ± 0.00121	0.00360 ± 0.01646	0.03622 ± 0.00167	0.013	70.8	299.5 ± 4.3	80.7
1055	1123.9	63.8 ± 0.1	0.01787 ± 0.00388	0.00461 ± 0.02962	0.06130 ± 0.00205	0.017	93.2	304.8 ± 4.8	71.6
1130	377.1	69.8 ± 0.1	0.02801 ± 0.00591	0.02533 ± 0.02013	0.08913 ± 0.00095	0.091	100.0	291.2 ± 3.4	62.3

¹ J—characteristic of the neutron flux during irradiation of samples; *—radiogenic ⁴⁰Ar.

, Figure 5), an age plateau is distinguished. In the high-temperature part of the spectra of biotites X-1041 (leucogranite of the Voylochev massif), X-1042 (granite of the Chernovin massif), and X-1047 (granodiorite of the Kurchum massif), a plateau of three or more successive stages is observed, the proportion of which is less than the accepted value of 60% of the ³⁹Ar released.

A summary of the thermochronological data we obtained, as well as those obtained earlier, is shown in Figure 2 and in Table 1.

5. Discussion

5.1. Ar/³⁹Ar Dating

In cases where a high-temperature plateau of three or more successive stages is observed, and the proportion of which is less than the accepted value of 60% of the ³⁹Ar released (biotite samples X-1041, X-1042, X-1047), the value of the high-temperature plateau corresponds to the age of closure of the isotope system of the mineral, and the presence of a plateau in the low-temperature part indicates a complex thermal history of the rock. This may be due to prolonged cooling or superimposed heating. The reason for this heating could be the belts of “postbatolite” dikes (age 286–267 Ma) attributed to the Mirolyubov complex (Figure 2) [5].

When comparing the obtained and published age data for the granitoid complexes of the Kalba–Narym batholith and the Chechek metamorphic complex, it can be noted (Table 1, Figure 2) that for each of the studied objects, the measured age correlates with the stability of isotope systems, decreasing in the series U/Pb for zircon → ⁴⁰Ar/³⁹Ar for muscovite → ⁴⁰Ar/³⁹Ar for biotite → ⁴⁰Ar/³⁹Ar for feldspar → AFT dating. This sequence is an independent confirmation of the assumption that the ⁴⁰Ar/³⁹Ar datings obtained correspond to the closing time of their K/Ar isotope system [55].

5.2. Thermochronology

The totality of all the dates measured by us and taken from the literature, shown in Table 1, is used in the thermochronological diagram (Figure 6). Each dating corresponds to the value of the closing temperature of the isotope system (T_c) or the formation (T_f) of the rocks as a whole in the case of U/Pb dating by zircon.

The formation of metamorphic rocks of the Chechek structure was simultaneous with the introduction of the Gabbro–Surov massif (see the geological description in part 2 of the article) [61]. This is confirmed by isotope dating data (Table 1, Figure 2). For muscovite from the Chechek metamorphic complex, an age of 312 ± 3 Ma was obtained by the ⁴⁰Ar/³⁹Ar method [65], consistent with the age of gabbro (313 ± 1 Ma [61]). At the same time, estimates of the conditions of metamorphism of samples taken at various sites of the structure are T = 665–720 °C, P = 4–6 kbar [65]. This corresponds to depths of 13–20 km, with an average of 16 km.

A close age (304 ± 4 Ma) was obtained for muscovite from the tectonic mélange zone at the southwestern contact of the rocks of the Surov gabbro massif (Figure 2a). The closure of the biotite isotope system from the same sample at 289 ± 3 Ma occurred 15 Myr later. To explain the sequential closure of isotopic systems of rocks formed at a depth of 13–20 km, one assumption about the establishment of thermal equilibrium is not enough. It is logical to assume that after its formation, the rocks of the Chechek structure experienced exhumation with an amplitude of several kilometers in the process of moving along the thrust (Figure 2 and Figure 3). Assuming that the amplitude was ~5 km in 23 Myr, this corresponds to a velocity of more than 0.2 mm/year, which confirms the tectonic nature of exhumation.

Indeed, a fairly rapid rise with an erosion rate significantly exceeding 0.33 mm/year could be caused by intense tectonic processes, as we assume, associated with the formation of a cover-thrust structure. For example, a study of the denudation rate of the Longmen Shan Ridge [73], located on the eastern edge of the Tibetan Plateau and characterized by steep terrain and a rate of horizontal contraction of up to 3 mm/year, derived denudation rates from 0.15 to 0.5 mm/year. The highest rates of exhumation and denudation are localized in the hanging walls of large thrusts (Figures 2 and 8 in [73]), emphasizing the role of tectonic structures in regulating the nature of denudation and topography throughout the Longmen Shan range.

In the history of the Chechek structure, after the closure of the biotite isotope system (289 ± 3 Ma), there is a break of 200 Ma (Table 1, Figure 2 and Figure 6), which ended with the closure of the AFT system from granite gneiss 87 ± 6 Ma ago [17]. Similar values of the AFT age are also observed in the southern extension of the ISZ for granitoid samples—72 ± 4 and 94 ± 4 Ma (Figure 2) [17].

The granitoid massifs of the Kalgutin, Kalba, and Monastery complexes were formed in the early Permian (297–286 Ma). Closure of the biotite isotope system in them (Chernovin, Asubulak, Narym, Sebin massifs, Figure 2a and Figure 6) occurred in a narrow age range of 289–279 Ma. For individual samples, an intermediate plateau with an age of 261–265 Ma is recorded in the low-temperature part of the spectrum. As noted above, this can be explained by the thermal effect of the dikes of the Mirolyubov complex on the isotope system of biotite.

Thus, it can be noted that for the massifs of granitoid complexes of the Kalba terrane, which were formed in the range of 308–284 Ma, the closure of the biotite isotope system occurs in a narrow age range—289–279 Ma. This can be explained by thermal equilibration after the formation of the massifs occurred at a shallow depth of 5–12 km. Although, given the elongated shape of the granitoid massifs in the north-west direction and the wide manifestation of thrust structures (Figure 2), it is possible to assume the closure of the biotite isotope system during tectonic exhumation.

⁴⁰Ar/³⁹Ar dating by feldspar (246 ± 3 Ma) from leucogranite in the central part of the Sebin massif is 33 Ma younger, and AFT dating (229 ± 21 Ma [70]) is 50 Ma younger than the age of massif formation.

The observed sequence of closure of low-temperature isotope systems for samples of the Sebin massif is radically different from the sequence observed for the Chechek structure (Figure 6). For the latter, the closure of the AFT system occurred 87 ± 6 Ma ago, 226 Ma after its formation, which may well be explained by gradual exhumation due to erosive denudation. At the same time, the closure of the AFT system in the 3 samples of leukogranites of the Sebin massif occurred much earlier. We believe that the tectonic nature of the exhumation of the Sebin massif should be used to explain this difference. We assume thrust motion as a probable exhumation mechanism (Figure 2b).

5.3. Geological Implications

The evolution of the CAFB includes a long history of the formation of late Neoproterozoic–Paleozoic orogens along the border of the Siberian continent. In the late Paleozoic, the closure of oceanic basins and the collision of the Eastern European and Siberian continents occurred. The process of tectonic transition from subduction to collision in this vast territory, characterized by the formation of cover-thrust and shear structures, is still poorly understood, with the exception, perhaps, of the Chinese Altai Orogen.

The Chinese Altai orogen is a continuation to the south of the Gorny–Altai terrane. The East-West Junggarian terrane is a continuation to the south of the Kalba–Narym terrane. The southeastern extension of the ISZ is the border between the Chinese Altai orogen and the East-West Junggarian terrane (Figure 1). In the Chinese Altai orogen, a variety of publications [34,36,74] prove the late Paleozoic transition from subduction to late Carboniferous collisional high-temperature metamorphism and early Permian granite magmatism [75,76]. It was revealed [36] that in the late Carboniferous (322–295 Ma) there was a thickening of the orogen, the formation of thrust and high-temperature collision metamorphism, including the ISZ, associated with the collision of the Chinese Altai orogen with the East Junggarian terrane within the NE-SW compression. Then there was a stretching of the orogen (295–283 Ma), which could be responsible for the high thermal gradient and the development of widespread early Permian magmatism. It was followed by a transpression event (folding and lateral strike-slip) (283–253 Ma) associated with the oblique convergence of the Chinese Altai orogen with the Eastern Junggarian terrane through the ISZ. It is recorded by ⁴⁰Ar/³⁹Ar dating of syn-strike-slip hornblende and biotite [35,77].

For the AACS, we have obtained new geological and geochronological data that support a scenario similar to the Chinese Altai orogen and well support the manifestation of late Carboniferous–early Permian cover-thrust tectonics and related manifestations of magmatism (Kalba–Narym batholith) and metamorphism within the ISZ. It is shown that in the area of Ust’-Kamenogorsk, the late Carboniferous Surov gabbro complex and Chechek metamorphic rocks are located in a cover-thrust structure deformed into dome-shaped folds. It was formed as a result of the general NE-SW compression, which is also characteristic of the Chinese Altai. Metamorphic rocks form a tectonic mélange by the framing of a large hot gabbro body embedded in the thrust structure. Tectonically higher are the metamorphic rocks of the green schist facies of metamorphism, formed by middle Devonian sedimentary rocks (Kystav–Kurchum Fms.). They are characterized by shallow recumbent folds with axial planes and mineral linearity sinking to the northeast and east at angles of 35–10°, which also indicates a regional NE–SW compression of the region. As a result of the formation of the cover-thrust structure, volcanogenic sedimentary and sedimentary rocks of the AACS are everywhere compressed, close to isoclinal folds of the northwestern strike of axial planes inclined also to the northeast and east [2,10].

The ISZ is located between the late Silurian–early Carboniferous formations of the Altai active margin (Rudny–Altai terrane) and the Middle Devonian–early Carboniferous turbidites of the pre-arc trough (Kalba–Narym terrane) and probably represents a fragment of the transition zone from subduction to collision, limited by strike-slip. Lateral strike-slip displacements within the studied region are manifested in Z-shaped folds with vertically sinking hinges formed in narrow fault zones represented by green schist, mylonites, and blastomylonites from rocks of the Chechek metamorphic complex of the epidote-amphibolite facies of metamorphism. Lateral strike-slip displacements are most fully developed on the border of the ISZ with the Rudny–Altai terrane and are an object for further detailed study.

The thermochronological reconstruction (Figure 6) carried out for the Chechek metamorphic rocks and the Kalba–Narym batholith indicates the manifestation of a large late Carboniferous–Middle Triassic (312–229 Ma) tectonic stage, as a result of which the rocks were eroded and brought to the surface. Thus, the Chechek metamorphic rocks characterizing the zone of complete tectonic mélange in the late Carboniferous–early Permian (312–289 Ma) were brought to the surface at a denudation rate of >0.2 mm/year. The most complete set of dates obtained by various methods characterizes the leucogranites of the Sebin massif of the monastery complex of the Kalba–Narym batholith. It was found that if based on the results of ⁴⁰Ar/³⁹Ar dating of feldspar and AFT dating, in the range of at least 250–230 Ma years ago, the leucogranites of the massif experienced an tectonic exhumation during thrust motion.

6. Conclusions

As a result of new geological and geochronological data with thermochronological interpretation, the sequence of the AACS formation stages in the Kalba–Narym terrane has been established:

• At an early stage in the Carboniferous–early Permian (312–289 Ma), the ISZ formed as a shallow thrust structure under conditions of NE–SW compression. Synchronously with the thrust, the intrusion of the gabbro of the Irtysh complex (Surov massif and others) and the formation of tectonic mélange with gabbro-cataclasites and metamorphic rocks occurred. The foundation of the Chechek dome structure corresponds to this stage.
• At the next stage, the formation of a cover-thrust structure involving turbidites of the Middle Devonian–early Carboniferous (Kalba–Narym terrane) led to melting at the middle levels of the thickened crust and the formation of the early Permian Kalba–Narym batholith (297–284 Ma).
• Rapid denudation of the created orogen occurred at the final stage of collapse until the Early Triassic (279–229 Ma).

Thus, in contrast to the currently prevailing ideas about the significant role of strike-slip tectonics, thrust and cover-thrust processes played a key role at almost all stages of the formation and evolution of the AACS. The Irtysh Shear Zone is located between the late Silurian–early Carboniferous formations of the Altai active margin (Rudny–Altai terrane) and the Middle Devonian–early Carboniferous turbidites of the pre-arc trough (Kalba–Narym terrane) and is probably a result of the subduction zone processes, followed by collisional processes, and only then—strike-slip processes.