Mineralogy, Geochemistry, and Geochronology of the Yehe-Shigna Ophiolitic Massif, Tuva-Mongolian Microcontinent, Southern Siberia: Evidence for a Back-Arc Origin and Geodynamic Implications

Abstract

The new results have been represented of mineralogical–geochemical and geochronological studies of rocks of the Yehe-Shigna ophiolite massif located in the Tuva-Mongolian microcontinent in the northern part of the Central Asian orogenic belt (Eastern Sayan, Southern Siberia). The Yehe-Shigna ophiolite massif is part of the Belsk-Dugda ophiolite belt. The structural position, age, and geochemical characteristics of the belt indicate its formation in the setting of the back-arc basin of the Shishkhid intraoceanic island arc, developing in the period of 810–750 million years. It is assumed that together with the same-age formations of the Oka accretion wedge and the Sarkhoi active margin, it formed on the convergent margin of the Gondwana supercontinent. Its basement is represented by the Archean-Early Precambrian crystalline rocks and carbonate cover (“Gargan Glyba”). The gold-bearing Neoproterozoic deposits with dominant gold-telluride assemblages are localization in large ophiolites thrust zones along with the frame of the “Gargan Glyba”. They are allochthonous with respect to the Late Neoproterozoic-Cambrian Tuva-Mongolian island arc of the Siberian continent. A similar type of gold deposit is probably worth looking for ophiolites thrust zones in other Precambrian Gondwana-derived microcontinents.

1. Introduction

A large composite Tuva-Mongolian microcontinent (massif) (1000 × 600 km) is located in the structure of the Central Asian orogenic belt along the southern margin of the Siberian craton (Figure 1). Its northern part (Figure 2) within the Eastern Sayan is represented by Early Precambrian crystalline basement and carbonate cover (“Gargan Glyba”), fragments of Early Neoproterozoic Dunzhugur ocean island-arc and Late Neoproterozoic Shishkhid ocean island-arc, Oka accretion prism, and Sarkhoi continental arc. They are jointly overlain by the Ediacaran-Cambrian carbonate cover up to 4 km thick [1,2,3,4,5,6,7,8,9,10,11,12,13,14].

The geodynamic nature of the Tuva-Mongolian continental block is a challenging problem: many similar Precambrian fragments are dispersed throughout the Central Asian orogenic belt, and at least two diametrically opposite hypotheses are discussed for their origin. The first point of view reconstructs the Neoproterozoic Tuva-Mongolian microcontinent as a sliver of the continental crust of the Siberian craton [14,15,16,17]. The second postulates that all terranes are fragments of East Gondwana and drifted away from it from the late Neoproterozoic [18,19,20,21,22,23]. According to [21,22,23], the Late Neoproterozoic-Cambrian southwest margin of the Siberian continent has a basement that was formed during the subduction of the Paleoasian oceanic plate, comprising the collage of Precambrian Gondwana-derived microcontinents (Tuva-Mongolian, Dzabhan, Muya, and others), beneath the Late Neoproterozoic-Cambrian Tuva-Mongolian island arc.

The formations of the Ediacaran-Cambrian island arc are located along the margin of the Tuva-Mongolian microcontinent from the west and east, the interaction with which are intensively interfered by Late Paleozoic deformations (Figure 2). From the north, through the shift of the Late Paleozoic Main Sayan Fault, these structures border on the Siberian Craton.

Ophiolite belts localized in the thrust structures of various ages are widely distributed along the frame of the “Gargan Glyba” and in the Oka accretion prism [2,9,14]. The most recent is the Late Paleozoic, whose role in the formation of the structure of the Eastern Sayan is still poorly evaluated. The formation of large deposits of ore gold (Zun-Holba, Vladimirskoye, Pioner, Zun-Ospa, etc.) is spatially related to the thrust ophiolite structures [24,25,26].

The gold-bearing deposits are associated with either Neoproterozoic or Early Paleozoic granitoids, belonging to the two main orogenic stages of the East Sayan geodynamic evolution. At ~850 Ma, during the Neoproterozoic stage, deposits with dominant gold-telluride assemblages formed in association with granitoids characterized by geochemical features of island arc granites. During 458–439 Ma, in the Early Paleozoic stage, gold–tetradymite, gold–stibnite, gold–telluride, and gold–bismuth–sulfosalt assemblages were formed that are spatially associated with orogenic granites with different geochemical compositions. Most gold-bearing mineral assemblages are intersected by post-mineral Late Paleozoic dykes. The origin of these different gold–sulfidetelluride assemblages is explained, with their genetic association with granitoid intrusions of different ages and compositions [26].

The Neoproterozoic gold-bearing deposits are located in large thrust zones along the frame of the “Gargan Glyba”, in which ophiolites are widely distributed (Figure 2). The northern Hamsarin and southern Ilichir zones are distinguished, which connect in an easterly direction in the upper reaches of the river Onot, where they form a large ophiolite nappe, represented by the Ospa basite-ultramafic massif. The northern thrust zone is composed of the Dunjugur island arc-ophiolite, and the southern thrust zone of oceanic serpentinized harzburgites, dunites, gabbro-pyroxenites, cumulative and stratified gabbro.

The ophiolote are thrust over the formations of the “Gargan Glyba”, serpentinite melange, and olistostrome tectonic plates developed at their base [1,2,3,4,9,14]. To the north of the “Gargan Glyba” on the border of the Oka accretion prism and the back-arc basin of the Shishkid island arc, there are fragments of Late Neoproterozoic ophiolites of the Belsk-Dugda belt. This belt includes one of the largest Yehe-Shigna massif (4 × 5 km) (Figure 3), located near the Siberian craton.

The formation of the Shishkhid ophiolite is associated with intraoceanic subduction, which took place during 805–600 MA [7,9].

This paper presents new mineralogical, geochemical, and geochronological characteristics of the Yehe-Shigna ophiolite massif, which characterize its geodynamic nature and structural position. Based on the obtained results, the presented interpretation of the geodynamic situation of the formation of structural complexes of the Tuva-Mongolian microcontinent in the Late Neoproterozoic is important for metallogenic analysis.

2. Materials and Methods

2.1. Chemical and Mineralogical Methods

In total, 45 samples of the Yehe-Shigna ophiolite and surrounding volcanic-sedimentary rocks were studied (Supplementary Material Table S1). The analyses of the major, trace, and rare earth element compositions were carried out at the Analytical Center for Multi-Elemental and Isotope research (VS Sobolev Institute of Geology and Mineralogy, Novosibirsk, Russia). Mineral chemistry was determined by wavelength-dispersive analysis using electron probe microanalyses (JEOL JXA-8100) and MIRA 3 LMU scanning electron microscope, with an attached INCA Energy 450 XMax at the Sobolev Institute of Geology and Mineralogy, Russian Academy of Science, Novosibirsk, Russia (Analytical Center for Multi-Elemental and Isotope Research, SB RAS). The accelerating voltage was 20 kV, the probe current was 50 nA, the beam size was 3–5 μm, and the signal accumulation time was 10 s. The standards used were natural and synthetic silicates and oxides. The detection limit for oxides was 0.01–0.05 wt.%. The major element composition was determined using VRA-20R X-ray fluorescence. The analytical errors were generally less than 5%. Trace elements (including rare-earth elements) were analyzed in solutions by inductively coupled plasma mass spectrometry (ICP-MS) using a Finnigan Element mass spectrometer [27]. The detection limits for trace elements were in the range of 0.01–0.2 μg/L. Microtextural observations of rocks were performed by means of light microscopy (AxioSkope.A1 Zeiss).

2.2. U–Pb Age Determination

U–Pb zircon dating was done by LA-ICP-MS method at Geological Institute, Siberian Branch of the Russian Academy of Sciences (Ulan-Ude) on an Element XR mass spectrometer with a UP-213 laser attachment [28]. Zircons were irradiated for 30 s by a laser beam with a frequency of 10 Hz and diameter of 30 μm. Evaporated particles from the laser attachment were transported to the mass spectrometer by a pure helium flow. The signal drift of the measured isotopes was corrected, background signals accounted for, and isotope ratios and their errors calculated in the Glitter program [29]. The external standard was 91,500; Plešovice and GJ-1 zircons were used as the controls. The relative error in measuring the isotope ratios in the standard zircons varied within 1–3%. The relative error of the weight-average concordant age of the control zircons was no more than 2% of their certified age. The Isoplot-3 program [30] was used to calculate age values by constructing a concordia plot. Four isotope ratios were measured: ²⁰⁷Pb/²⁰⁶Pb, ²⁰⁶Pb/²³⁸U, ²⁰⁷Pb/²³⁵U, and ²⁰⁸Pb/²³²Th. The errors for single analyses (ages) are presented at a level of 1σ; the errors in calculating the concordant ages and recalculating with concordia, at 2σ.

2.3. ⁴⁰Ar/³⁹Ar Dating

Monomineral fractions were extracted using standard magnetic and density separation methods. ⁴⁰Ar/³⁹Ar isotope studies [31] with gradual heating were carried out at the Collective Use Center for Multielement and Isotope Studies at the Sobolev Institute of Geology and Mineralogy, Siberian Branch of the Russian Academy of Sciences (Novosibirsk). Sample charges, together with MCA-11 muscovite (age 311.0 ± 1.5 Ma), were used as a monitor and calibrated with international standard samples of biotite LP-6 and muscovite Bern-4M after [32], were wrapped in aluminum foil, placed in a quartz ampoule, and vacuum sealed after the air was pumped out. Then, the samples were irradiated in the cadmium channel of a VVR-K scientific reactor at the Physical–Technical Institute of Tomsk Polytechnic University (Tomsk, Russia). The neutron flux gradient did not exceed 0.5% of the sample size. The experiments with gradual heating were carried out in a quartz reactor with an external heater. A dry experiment on determining ⁴⁰Ar (10 min at 1200 °C) did not exceed 5 × 10–10 ncm³. Argon was purified with ZrAl-SAES getters. The isotopic composition of argon was measured on a Noble Gas 5400 mass-spectrometer (Micromass, Manchester, England, UK). The analytical measurement errors correspond to the ±1σ interval. To correct the argon isotopes mixing that formed during irradiation by Ca, Cl, and K, the following coefficients were used: (³⁹Ar/³⁷Ar)_Ca = 0.000891 ± 0.000003, (³⁶Ar/³⁷Ar)_Ca = 0.000446 ± 0.000004, and (⁴⁰Ar/³⁹Ar)_K = 0.089 ± 0.001. Prior to measurements, the samples were preliminarily degassed at a temperature of 350 °C. To control isotope discrimination of the mass spectrometer, portions of purified atmospheric argon were measured. The average value of the ⁴⁰Ar/³⁶Ar ratio for the measurement period was 299 ± 1.

3. Geology, Mineralogy, Geochemistry, and Geochronology of the Yehe–Shigna Ophiolite Massif

3.1. Geology Setting

The Yehe-Shigna ophiolite massif is located in the East Sayan, in the upper reaches of the Yehe-Shigna River, in the interfluve of the r. Urik and Bolshaya Belaya (Figure 1 and Figure 2). It forms an oval elongated in the western direction with an area of about 20 km² (Figure 3). The most common rocks are serpentinites, serpentinized ultramafic rock (dunites and peridotites), and gabbro; smaller pyroxenite bodies are less common. The ophiolite massif and, in general, the Belsk-Dugda ophiolite belt are located in the fault zone, possible Late Paleozoic, between the Ediacaran-Cambrian carbonate rocks of the Tuva-Mongolian microcontinent cover and the formations of the Oka prism (Figure 2).

The ophiolite massif is located in the thrust structure. The Ediacaran-Cambrian carbonate rocks are thrust over them. In turn, ophiolites are thrust onto the black shale strata containing layers of sulfide ores. Carbonaceous formations in it are represented by powerful bundles of schist and rhythmically layered sand-shale-siliceous rocks, turning into an olistostroma with angular fragments, semi-rolled and rounded blocks of limestone, siliceous rocks, and carbonaceous–siliceous shales. In the composition of this terrigenous sequence, interlayers of metavolcanites rocks are also observed, represented by altered rhyolites, rhyodacites, and andesites. As they approach the ultrabasic rocks, carbonaceous deposits acquire a clastic texture and are gradually replaced by greenish calcinated schist.

The ophiolite massif composited relatively slightly of modified dunites and harzburgites lying on serpentinite schist (Figure 3). Harzburgites have been modified partly to serpentine–pyroxene and serpentine–amphibole rocks. There are areas of talc-carbonate and chlorites rocks in serpentinite shales. Gabbro composes dikes among ultramafic rocks and reaches a thickness from several tens’ meters to half km across. Gabbroids are mainly represented by amphibole gabbro of different granularity and gabbro-diorites. The contacts of gabbroid dikes and ultramafic are zones of crushing, composed of talc-carbonate chlorite–zoisite–amphibole rocks.

3.2. Mineralogy

Serpentinized dunites and serpentinites of antigorite, less often lizardite and lizardite–antigorite compositions predominate in various degrees; serpentine–pyroxene, serpentine–clinopyroxene, and serpentine–amphibole rocks are present in smaller quantities. Primary silicate minerals are present in relics: olivine corresponds to forsterite, and clinopyroxenes to augite and diopside. There is also magnesian hornblende, the appearance of which is apparently due to the amphibolization of pyroxenes [33].

In the fault zones, serpentinites have been transformed into talc-carbonate rocks, small chlorite bodies, and carbonization zones are located in places among serpentinites. The chemical composition of ultramafic rocks corresponds to mantle restites—they are enriched with Mg, Cr, and Ni and depleted in Si, Al, Ti, and incoherent elements [33].

According to the chemical composition, two groups of accessory chromespinelides are distinguished. Chromium spinel of the first group (Shp I) is characterized by a low content of Al²O³ (5–8 wt.%), Fe²O³ (0.2–3.0 wt.%), and MgO (6–9.8 wt.%), with a high value of Cr# (0.83–0.89) and Cr2O3 (61–65 wt.%) (Table S2). In the classification diagram [34], chromespinelides of the first group correspond to chromite (Figure 4). In some cases, chromite is partially or completely replaced by chromium-magnetite and magnetite containing small relics of chromium spinels. Two types of magnetite (Mgt) are present in serpentinized ultramafic rocks: Mgt I contains small impurities of Al₂O₃, MgO, and ZnO; Mgt II practically does not contain Cr, Al, and Mn but is characterized by elevated concentrations of NiO (0.4–0.7) wt.%. Usually, magnetite with such a composition is formed during the serpentinization of ultramafic rocks. Accessory chromium spinelides of the second group (Shp II) have a composition: Al₂O₃ (10–15 wt.%,) and Fe₂O₃ (4–8 wt.%), relatively low chrome (Cr# = 0.62–0.8), wide variations in FeO and MgO contents.

Chrome spinel of the second group on the classification diagram corresponds to subferrichromite and subferrialumochromite (Figure 4). Microinclusions of chloridized clinopyroxene and amphibole of isometric and irregular shape are present in chromium spinel.

When approaching the fracture sites in which chromite is replaced by chrommagnetite and magnetite, FeO increases from 18–22 wt.% to 26–29 wt.%, MgO decreases accordingly from 7–9 wt.% to 1.3–6.5 wt.%, and MnO increases to 1.2 wt.% and ZnO up to 0.6–0.9 wt.%, with a constant content of Fe₂O₃ (4–8 wt.%). Chrommagnetite is characterized by the presence of TiO₂ (0.26–0.7 wt.%), which is absent in chrommagnetites of the first group (Table S2). Chrommagentite and magnetite make up the edges in the grains of chromium spinel.

In addition to chrome spinel, ilmenite FeTiO₃ is present; in composition, it is close to ilmenites from ultrabasic and basic rocks, except for increased content of MnO-5 wt.%, which is characteristic of granitoids, alkaline ultramafic, and carbonatites (Table S2).

In the discrimination diagram Al²O³–Cr²O³, the points of the compositions of the chrome spinel of the Yehe-Shigna massif fall into the field of ophiolite chromites (Figure 5a). According to the ratio Al₂O₃–Fe²⁺/Fe³⁺, the studied chrome spinel corresponds to the compositions of chrome spinel from suprasubduction peridotites (Figure 5b). In the Mg#–Cr# diagram, chrome spinel shows some variation in Mg# values, but, nevertheless, most of the points of chromespinelide compositions fall into the field of post-arc and pre-arc ophiolites (Figure 5b), while chrome spinel of the Ospa-Kitoy massif is localized near the boninite field.

Gabbro is represented by several varieties, among which hornblende-reached gabbro and gabbro-diorites predominate. Rocks are altered in varying degrees. There are several textural varieties of gabbro-rocks of massive texture with a relatively uniform distribution of aggregates of isometric grains of dark-colored minerals prevail (Figure 6a,c); there are also areas with breccia-like, lenticular, and lenticular-striped textures, the appearance of which is due to tectonic deformations (Figure 6d,e).

Coarse-grained gabbro is composed of large grains (up to 1 cm across) of hornblende, composing up to 40 vol.% of the rock (Figure 6b). Here, the primary amphibole is least susceptible to secondary changes; it has a clear pleochroism, from brownish to greenish colors. Thin borders of chlorite–clinozoisite and actinolite–tremolite aggregates are occasionally found along the edges of large hornblende secretions. Numerous poikilitic inclusions of plagioclase in large amphibole grains are marked. The chemical composition of the primary amphibole corresponds to magnesium hornblende with increased content of TiO₂ up to 2.39 wt.% (Table S1). In the finer-grained bulk, gabbro is represented by an aggregate of epidote–chlorite–clinozoisite–albite composition with veined separations of titanite, clinozoisite, and rare scales of muscovite. Albite is often sericitized in the mass of secondary minerals. Rock structures in thin sections are porphyritic, hypidiomorphic-grained, poikilitic, and granoblastic.

Gabbro-diorites are medium-grained massive rocks with a relatively lower content of mafic minerals. The rock is composed of magnesium hornblende, partially replaced by amphibole of the tremolite–actinolite series (Table S3). The rock is composed of plagioclase–clinozoisite aggregate, with an admixture of chlorite and epidote. Plagioclases are sericitized, form large hypidiomorphic crystals, rarely skeletal grains among secondary minerals, and are partially subject to sericitization. Clinozoisite composes relatively plate-shaped crystals or veined aggregates. Primary plagioclases correspond in composition to andesine and oligoclase (Table S4). They are found most often in the form of small relics among aggregates of secondary minerals—albite, epidote, quartz, and amphiboles from the actinolite–tremolite series.

The varieties of gabbro are represented by rocks of chlorite–clinozoisite–tremolite–actinolite structure, composed of complete pseudomorphoses of tremolite–actinolite along the hornblende, localized among the aggregate of isometric clinozoisite grains, scaly chlorite aggregates with relicts, or skeletal albite grains. The structure of rocks is porphyroblaste and granoblaste. Metavolcanic rocks are represented by light gray color, spotted or striped texture, and porphyritic structure (Figure 6e). They are composed of quartz, feldspar, and biotite. Feldspar and biotite may be changed in the sericitization of plagioclases and chloritization of biotite. Porphyry inclusions are represented by biotite and quartz (Figure 7d).

3.3. Geochemistry

Gabbroid form a range of melanocratic gabbro to gabbro-diorites and quartz diorites, where SiO2 contents vary from 44.40 to 69.14 wt.% (Table S1). On Harker binary diagrams, gabbro and gabbro-diorites form uniform trends (Figure 8).

In Figure 8, the figurative points of serpentine–pyroxene rock composition are with Mg and Al content. With an increase in SiO₂, Fe, Mn, Ca, and Mg decrease, and the total alkalinity increases. There is no such dependence of Ti and Al. Pegmatoid gabbro enriched with large hornblende grains is characterized by a sharp enrichment of TiO₂, the contents of which reach 1.76 wt.%, whereas in the gabbro, they vary within 0.1–0.64 wt.%. The La-SiO₂ diagram shows that the distribution of elements in the studied rocks most corresponds to the trend of partial melting (shown by a solid line) and differs from the trend of fractional crystallization (shown by a dotted line).

The sum of rare earth elements in gabbro is low, ranging from 4.21 to 16.69 g/t. The studied mafic rocks have two types of REE distribution patterns (Figure 9).

In the first case, these are flat patterns without Eu anomaly (Eu/Eu* = 0.97–1.11), which are similar in configuration to the REE distribution patterns in the Izu-Bonin Island Arc basalts. La_n/Yb_n in rocks of this group varies from 0.6 to 2.3. The same group includes REE patterns in pegmatoid gabbro enriched with hornblende (sample XB-72). They are characterized by a relative depletion of light REE and a high content of heavy REE, resulting in a minimum La_n/Yb_n ratio (0.6). The second group is characterized by some enrichment with light REE and a positive Eu anomaly (Eu/Eu*-up to 2.8). In the rocks of this group, the La_n/Yb_n ratios have values from 1.6 to 3.4. The spider diagrams also show the difference in the geochemical characteristics of gabbro in the distribution of Th, Nb, and Ba. Some of the samples show positive anomalies and some negative.

In general, the patterns of incoherent elements in all studied gabbro are similar. All rocks are characterized by the presence of a positive anomaly in Pb, Sr, and Ti, as well as increased contents of large-ion lithophilic elements (LILE) and reduced contents, relative to N-MORB, of high field strength elements (HFSE). Most of the analyzed samples show a negative Nb anomaly (Figure 9a). The listed geochemical characteristics of the studied gabbro correspond to suprasubduction rocks formed in the back arc basin [44]. The metavolcanites correspond in chemical composition to low-alkaline andesites, rhyolites, and rhyodacites. SiO₂ contents vary from 60.99 to 71.93 wt.% (see Table S1). The total alkalinity of rocks varies in the range of 4.35–6.76 wt.%. The REE patterns show some enrichment of light REE relative to heavy ones (La_n/Yb_n ratio = 5.7–10.3) and a weakly pronounced negative Eu anomaly (Eu/Eu* = 0.54–1.12) (Figure 9b). The REE patterns for metavolcanites (Figure 9c,d) are close to those for gabbroids (Figure 9a,b). They have a negative slope, indicating the enrichment of rocks with incoherent elements relative to the basalts MORB and OIB.

This trend REE pattern corresponds to the direction of these relations from the incipient back-arc basalts (the field with the number 2 in Figure 10: Izu-Bonin-Marian pools) to the back-arc pools [44]. Thus, the composition of the gabbro of the Yehe-Shigna ophiolite massif and associated metavolcanites correspond to the compositions of rocks from back-arc ophiolites [45].

Figure 9. Chondrite-normalized REE diagram for gabbro (a,b) and metavolcanites (c,d). Data for the average chondrite, primitive mantle, enriched MORB (E-MORB), normal MORB (N-MORB), and ocean island basalt (OIB) are from Sun and McDonough [46].

Figure 9. Chondrite-normalized REE diagram for gabbro (a,b) and metavolcanites (c,d). Data for the average chondrite, primitive mantle, enriched MORB (E-MORB), normal MORB (N-MORB), and ocean island basalt (OIB) are from Sun and McDonough [46].

The Th/Yb–Nb/Yb discrimination diagram (Figure 10) shows the genetic nature of gabbro since the behavior of Th and Nb differs sharply in subduction environments—Th is more mobile than Nb, and in addition, the Th/Nb ratio also depends on the source of melting basalts [47]. The compositions of oceanic basalts, the formation of which is not associated with subduction, form a linear trend of a sequential increase in the values of Th/Yb and Nb/Yb ratios (shown by a solid arrow in Figure 10).

The arrival of the slab component disrupts this sequence in the direction of increasing the values of Th/Yb ratios, as a result of which the points of the compositions of the gabbro of the Yehe-Shigna array and the associating effusions approach the compositions of E-MORB (dotted arrow in Figure 10).

Thus, amphibole gabbros of the Yehe-Shigna massif correspond in geochemical characteristics to typical ophiolite gabbros formed in a suprasubduction environment and correspond to back arc basin basalts. The association of basite-ultramafic with rocks of andesite-rhyodacite composition is a characteristic feature of ophiolites formed in the back arc spreading zones [48].

3.4. Isotope Geochronological Dating

⁴⁰Ar/³⁹Ar dating. It was performed in the Analytical Center of IGM SB RAS by the method of stepwise heating of the monomineral amphibole fraction according to the method published in [31]. A sample of a coarse-grained amphibole gabbro was selected for dating, from which a monofraction of dark green amphibole corresponding to the composition of hornblende was isolated under the binocular. The selection of the sample is due to minimal secondary changes in the hornblende and the absence of inclusions of other minerals, which can be seen in the thin sections (Figure 7a).

Relatively reduced values of the Ca/K ratio were recorded, which indicates the favor of the four temperature stages, indicating a significant contribution of the modified amphibole (Figure 11). Ratios of the order of 170–140 for two high-temperature stages are observed. In the age spectrum, after two low-temperature steps with sharply inflated age values, associated with the effect of loss of ³⁹Ar due to recoil energy, an upward staircase is observed. The last high-temperature stage is characterized by 53% of the isolated ³⁹Ar and an age value of 736 ± 10 Ma. Considering that it is at high temperatures, as a rule, that argon is released from the amphibole lattice, the age of 736 ± 10 Ma should be considered the best approximation of the closing age of the isotopic K/Ar system of the hornblende. Thus, the formation of the hornblende and, accordingly, the amphibole gabbro of the Yehe-Shigna massif occurred no younger than 736 ± 10 Ma. This age estimate differs from the age of the ophiolites of the Dunzhugur massif and is close to the age of the ophiolites of the Shishkhid belt [7].

U-Pb dating on zircons. To clarify the age of basalt rocks, the U-Pb dating of zircons was carried out. Several monofractions of zircons were isolated from gabbro rocks of the Yehe-Shigna massif. The age of the ophiolites, determined by the U-Pb method using zircons isolated from hornblende gabbro (Vs-55-99 sample), showed three concordant clusters corresponding to three age dates: 806.5 ± 9.4 million years, 501.1 ± 16 Ma, and 288.8 ± 4.5 Ma (Figure 12). For the most ancient population of zircons, CL images are shown (Figure 13 ), which show a clear igneous oscillation zoning. The U/Tr ratio is in the range 0.41–0.71.

It should be noted here that the ages of about 800 and about 500 Ma were obtained by dating the same grains, where the central part of the grain is of ancient age, and the marginal part is relatively young. The dating of the second sample of zircons from gabbro showed a U-Pb value of 687.5 ± 7.6 Ma (Figure 14).

4. Results and Discussion

4.1. Genesis of Basite-Ultramafic Massif

The attribution of the Yehe-Shigna basite-ultramafic massif to the ophiolite association was previously based mainly on the geological relationships of ultramafic and host lithocomplexes represented by the black shale strata with interlayers of volcanites, sandstones, and carbonate rocks, which were considered as the upper part of the ophiolite association [33]. However, the key rocks of the ophiolite association—the cumulative and parallel dike complexes—are not discovered in the Belsk-Dugda ophiolite belt.

The conducted studies of the geological position and composition of the rocks of the Yehe-Shigna massif made it possible to reliably substantiate its ophiolitic nature. In addition to ultramafic rocks, other members of the ophiolite association are also present in the massif: first of all, gabbro and gabbro-diorites, as well as pyroxenites; gabbro-diorites, dikes of medium and basic compositions, plagiogranites, the nature of which is reliably unknown, are found in smaller quantities.

The study of the composition of the gabbro, in this case, is more informative for determining the genetic nature of ophiolites since ultramafic rocks are subject to intense secondary modifications. The complete absence of pyroxenes and their relics in the studied gabbro is characteristic. The presence of amphibole indicates a high water saturation of the primary melt. The crystallization of magmatic amphibole in ultramafic occurs in subsolidus conditions, at temperatures of 870–1000 °C, due to the infiltration of the melt through peridotites (in this case, rocks of the mantle wedge) [49]. Rocks have geochemical labels of suprasubduction basites, which include high water saturation, low potassium content, increased LILE content, and reduced HFSE, which is fixed by positive and negative anomalies of the corresponding chemical elements on spider diagrams and Harker variation diagrams. The same “suprasubduction” characteristics are also manifested in the ophiolites of the Shishkhid belt [7]. The formation of suprasubduction ophiolites occurs during partial melting of the mantle wedge due to dehydration of the subducting plate, which is confirmed, in this case, by the nature of the La distribution in the binary diagram of La-SiO₂ (see Figure 8), where the distribution of rock compositions most corresponds to the trend of partial melting shown in the paper [31].

The ophiolitic nature of the Yehe-Shigna massif is confirmed by the compositions of accessory chromium spinel, typical for the field of back-arc and pre-arc ophiolites (Figure 5). The ophiolites of the Yehe-Shigna massif, according to geochemical characteristics, most correspond to the suprasubduction ones formed in the back-arc basin [44,45].

Thus, the results of U-Pb and Ar-Ar dating of the rocks of the Yehe-Shigna massif suggest that the beginning of the formation of the ophiolites of the Belsk-Dugda belt occurred about 805 million years ago. During the period from 805 to 687 million years, the Yehe-Shigna ophiolites gabbro underwent deformations epigenetic tectonic-thermal events, which is reflected in the presence of corresponding Ar–Ar (736 Ma) and U-Pb (687.5 Ma) dating (Figure 6). The geochemical characteristics of gabbro indicate their formation in the setting of a back-arc basin. According to the structural position and age of manifestation, this back-arc basin probably belongs to the back-arc basin of the Shishkhid island arc, developing in the period of 810–750 Ma [9].

4.2. Geodynamics of the Tuva-Mongolian Microcontinent in the Late Neoproterozoic

We accept the point of view that the Tuva-Mongolian microcontinent is a fragment of Gondwana [18,19,20,21,22,23]; then, the considered geodynamic complexes and the Neoproterozoic gold-telluride deposits of the Eastern Sayan can be interpreted as formations of its active margin (Figure 15). In this case, the Archean “Gargan Glyba” and the tectonic nappe of the Early Neoproterozoic Dunhzugur oceanic ophiolites, including gold deposits [13,14,24,25,26], are a fragment of Gondwana. Their cover consists of the Irkut Formation marbles at the bottom and the Ilchir Formation shales rocks above. The Irkut Formation consists of a thickly bedded dolomitic marble up to 600 m thick, in places stromatolitic, and contains recrystallized chert beds. A basal conglomerate contains pebbles of quartz and metamorphic rocks. The Ilchir Formation is composed of dark grey schist along with sporadic interbeds of sandstone and limestone. The sedimentary succession is interpreted as a passive-margin shelf deposit [9]. These rocks were intruded at 785 ± 11 Ma by the Sumsunur tonalite complex that defines the upper age limit of sedimentation. It is likely that the sedimentary cover is not older than the early Neoproterozoic. The U-Pb age of the Sumsunur tonalites in the Gargan Glyba is close in age and geochemical parameters to the Sarkhoi volcanites—782 ± 7 million years [10,11]. Based on these data, the comagmatism of Sumsunur granitoids with volcanogenic rocks of the Sarkhoi formation is established.

A number of geodynamic coeval complexes with the Sarkhoi active margin: the Oka accretion prism, the back-arc ophiolites of the Belsk-Dugda belt, and the Shishkid intraoceanic island arc can be considered as a fragment of the convergent boundary of the Gondwana supercontinent for the Neoproterozoic.

The Oka accretion wedge consists of a tectonic alternation of Neoproterozoic turbidites (carbonate-terrigenous, siliceous-carbonate-shale and terrigenous), MORB- and OIB-like oceanic basalts, and mélange with blueschists. The minimum U-Pb age of detritus zircons from turbidites is 786 ± 9 million years [14]. The age of the glaucophane-green shale strata was determined at 640 ± 20 million years (Rb-Sr isochron). There are also tectonic slices of weakly metamorphosed terrigenous sequences with interlayers of black carbonaceous siltstones with diabase sills and gabbro-diabases of the tholeiitic series of type N-MORB with an age of 736 ± 43 million years (Sm-Nd method) [9].

The Shishkhid ophiolite is a well-preserved 13 km thick mafic-ultramafic assemblage that comprises (from bottom to top) residual ultramafic rocks (~6 km), layered and isotropic gabbro (~4.5 km), sheeted dykes (up to 0.5 km), a bimodal assemblage of basalt and rhyolite (up to 0.7 km), and andesitic pyroclastic rocks (~2 km). The volcanic rocks are overlain by a 3 km thick sedimentary sequence showing progressive subsidence of the volcanic edifice after cessation of volcanism. The sedimentary unit is unconformably overlain by Ediacaran-Cambrian platform sediments. SHRIMP U-Pb dating of magmatic zircons from rhyolite of the lower volcanic unit has yielded a concordant ²⁰⁶Pb/²³⁸U age of 800 ± 2.6 Ma that is interpreted to reflect the time of magma crystallization. The tectonic setting for the Shishkhid ophiolite is inferred to be similar to that of the Izu-Bonin “back-arc knolls extensional zone”. The dating of zircons from tephra and volcanic sandstones showed a period of arc formation in the interval 800–750 Ma (Kuzmichev et al., 2015).

According to the characteristics listed above: accretion complexes with turbidites, glaucophane-green shale strata, fragments of MORB- and OIB-like oceanic basalts, back-arc basins, and oceanic arcs, the Late Proterozoic convergent boundary (Figure 15) is comparable to the modern Western Pacific active margin. For example, a similar geodynamic situation is currently implemented in the Solomon Sea area—the Woodlark Basin [50,51], in SW Iberia [52,53], and other regions.

The structure of the Belsk-Dugda ophiolite belt of the back-arc basin, the Shishkid oceanic ophiolites, and, in general, the Oka accretion prism is largely disrupted by Late Paleozoic deformations, which complicated the primary relationships of the Late Neoproterozoic period of their formation. It is likely that the Early Permian age of zircons from the gabbro studied by us reflects this tectonic event. Well-defined Late Paleozoic thrusts (Figure 2) form a knee-shaped structure and dip into the north (for the latitudinal segment) and west (for the meridional segment). These directions of immersion of young fault structures are recorded throughout the Tuva-Mongolian microcontinent and are shown in detailed geological sections [5,9].

The Late Paleozoic thrusts are intensively interfered to the structure of the Precam-brian Tuva-Mongolian microcontinent and the Paleozoic metamorphic, igneous, and sedimentary rock complexes developed within it. They are most fully recorded in the Tunka ridge to the east along the Oka thrust zone, in which the northern Khamsarin and south-ern Ilchir ophiolite belts connect (Figure 2). Here, according to the geochronological and geological-structural data was a prominent Late Paleozoic stage of thrusting took plase [54,55,56]. The formation of the Late Paleozoic thrust structure occurred simultaneously with the activation of shear displacements along the Main Sayan fault located on the border between the Siberian Craton and the Tuva-Mongolian microcontinent (Figure 2).

5. Conclusions

• The Yehe-Shigna ophiolite massif is composed of serpentinized ultramafic rocks and gabbro. Hornblende gabbro predominates; gabbro-diorites and diorites are also present. Rocks are subject to secondary changes to varying degrees, expressed in the appearance of amphibole of tremolite–actinolite composition, clinozoisite, epidote, and albite. According to geochemical characteristics, ophiolites correspond to suprasubduction ophiolites formed in the zone of back-arc spreading.
• Geochronological dating of U-Pb (by zircon) and Ar–Ar (by hornblende) methods showed the gabbro age (806.5 ± 9.4 Ma), and several values of superimposed processes (736 ± 10, 687.5 ± 7.6, 501.1 ± 16 and 288.8 ± 4.5 Ma). The large variation in the age values of zircons and amphibole is due to the influence of epigenetic tectonic-thermal events on ophiolite gabbro, reflecting the geodynamic evolution of the northern part of the Central Asian orogenic belt in the Paleozoic [57,58,59]. This issue requires further study.
• The Yehe-Shigna ophiolite massif is part of the Belsk-Dugda ophiolite belt and characterizes the Late Proterozoic back-arc basin of the Shishkid oceanic island arc. Together with the coeval formations of the Oka accretion wedge and the Sarkhoi active margin, they represent a fragment (Tuva-Mongolian microcontinent) of the convergent boundary of Gondwana.
• The gold-bearing Neoproterozoic deposits with dominant gold-telluride assemblages are localization in large ophiolites thrust zones along the frame of the “Gargan Glyba” are generally associated with Neoproterozoic granitoids of the Sumsunur tonalite complex of the Sarkhoi continental arc. They are allochthonous with respect to the Late Neoproterozoic–Cambrian Tuva–Mongolian island arc of the Siberian continent. A similar type of gold deposit is probably worth looking for ophiolites thrust zones in other Precambrian Gondwana-derived microcontinents, including granitoids of the Neoproterozoic active margin.