Formation and Evolution of the Sarmatia Earth’s Crust (East European Craton): Evidence from the Dnipro-Donets Paleorift

Abstract

The East European Craton tectonics was studied based on the reconstruction of the stress field of Sarmatia. The Sarmatia internal plate field is influenced by induction stress from the movements of the Arabia and Scythia plates and the North Atlantic Ridge spreading. Stress from the plate boundaries is transferred inward and absorbed by the movable belts. The role of movable belts in the evolution of the lithosphere is considered on the basis of the Dnieper-Donet Paleorift. At the Hercynian-Alpine collision stage, the rift basin experienced inversion uplift and folding. The change in deformation modes was initiated by displacement of the geodynamic axes by 15° counterclockwise during the epoch with a total of 60° in Phanerozoic. The rift structure was deformed by longitudinal bending on diagonal strike-slips of opposite kinematics and longitudinal elongation on one-sided strike-slips or along the diagonal strike-slips. The framework of the deformations is made up of diagonal sutures form the “Longitudinal Extrusion Orocline” and the “Transverse Spreading Fan”. Consequently the West Donets Covered Folded Region and the Donets Foldbelt were formed and the Paleorift has turned into a movable foldbelt.

1. Introduction

The tectonic evolution of the lithosphere and Earth’s crust is characterized by its stages and cyclic nature and occurs permanently in a variable stress field, favoring one of the elementary deformation modes of compression, tension, or horizontal shifting [1, 2, 3, 4]. The key factor in the structural-kinematic evolution of the continental crust is the reactivation of tectonic faults along regular directions of the Planetary fracture grid [2, 3]. During crust evolution, individual tectonic frameworks are formed, controlling various age deformation planes of the original structure. The movements of activated rock masses along deformation frameworks initiate the transformation from block-faulted tectonics to fold-dislocated due to internal redistribution of masses and deformation structure formation [4]. Faults are re-mobilized in accordance with the nature of the actual stress field. In another geodynamic environment, a new deformation mode is activated, according to which the structural-dynamic manifestation of faults in the crust, the direction, and kinematics of massif and block movements change [5]. The Planetary grid is the basis of the global deformation framework, whose dynamics and kinematics are renewed through fault re-mobilization in a variable stress field. The invariance of deformation regimes along established directions of the grid is conditioned by spatial-temporal inversions of stress field parameters [2, 3, 6].

Tectonic deformations disrupt the isotropy of the lithosphere by dividing it into tectonic blocks and plates through the formation of fractures, faults, linear zones of fracture-dislocations, and mobile and collision belts. The character of the divisibility is conditioned by the anisotropy of the lithosphere, temporal changes of Planetary geodynamic processes, and directions of the Planetary fracture network. Faults disrupt the consolidated crust through fracturing and influence the character of crust structural evolution due to mobility and the ability to remobilize [2, 3, 4, 5, 6]. Faults consist of several different age shear deformation zones of approximate azimuth orientation, so their structure is determined by the genetic types of fractures that compose them [3, 5]. Faults form deformation frames in the directions of the Planetary grid [2] during their evolution. The grid has a steady orientation in geochronology relative to the figure of the Earth [2, 3]. The horizontal axes of tension (σ1) and extension (σ3) are located according to modern meridians and parallels, but in earlier geological epochs were rotated at angles multiple to 15 [2, 3, 5]. The periodicity of stress field inversions is due to the change in Planetary rotation parameters, so the periodicity of changes in deformation regimes is close to the geological epoch [2, 3]. According to the new mode, the genetic type of faults, directions, and kinematics of movements in different age structural surfaces change [2, 6].

The internal plate field of the East-European Craton (EEC) is most influenced by induction stress from the movements of the Arabia and Scythia continental plates and the North Atlantic Ridge spreading [7, 8, 9, 10]. To understand the deformation mechanism in areas of lithosphere plate collision, the direction and kinematics of displacements of rock masses, located between the plates, are usually reconstructed. Two vertical stress outflow directions: pulling down the lithosphere edge by a descending convection current during subduction, or squeezing upwards with uplift and fold formation, have a vivid structural manifestation in the modern relief and are relatively easily determined. However, within continental plates, significant horizontal movements occur within collision belts that are decisive for structure formation. The structural results of lateral displacements in intra-plate collision belts are not always reflected at the surface and are less confidently distinguished based on the analysis of structural patterns of deformations [2, 3, 4, 5, 6, 7]. The result of deformations by the mechanism of lateral squeezing of masses along the belts is the strike-slip structures as a diagnostic indicator [4, 5, 7]. The division of plates into relatively independent micro-plates by movable belts contributes to geodynamic independence and increases the overall plate mobility. Intra-continental rift belts are movable zones capable of absorbing and accumulating stress. When studying the mechanisms of intra-plate deformations, the general kinematics of movements of distant collision belts are taken into account [7, 8, 9, 10].

The core of the EEC is composed of three segments: Fenno-Scandia, Sarmatia, and Volga-Uralia (Figure 1) [11]. The internal tension field and deformations of the Sarmatia crust are most influenced by induction stress from the movements of the Arabian and Scythian continental plates and the North Atlantic Ridge spreading [7, 8, 9, 10]. During the pre-rift stage of evolution, due to lithosphere anisotropy, the Sarmatia was divided into West and East micro-plates and smaller blocks by ancient collision suture and seams [12]. At the rift stage, through the formation of the Dnieper-Donets Paleorift (DDP), a new division of the Sarmatia lithosphere occurred, isolating the massifs of the Ukrainian Shield (USh) and Voronezh Anticlise (VA). During the Hercynian, Cimmerian, and Alpine platform activation, the rift basin underwent uplift, folded, and over-thrust deformations and transformed into a collision foldbelt [6, 8, 9, 10, 16]. The problems of the formation and evolution of the Earth’s crust of the EEC are discussed below on the basis of the latest data on tectonics and geodynamics of the evidence from Sarmatia and DDP.

Figure 1.

Position of the Dnieper-Donets Paleorift on the Sarmatia plate and the Eastern European Craton (inset), by [11, 12] with additions, by [13, 14, 15]. Symbols: 1—intersegment zones; 2—Paleorift boundaries; 3—faults: a—interplates, b—intermega-blocks, c—interblocks; 4—tectonic seam (H-SM-Kherson-Smolensk) and neo-tectonics strike-slip zones; 5—volcano belts (OMBP—Osnytsko-Mikashevychsky; SYVP—Stavropol-Yertil); 6—kinematic of displacement; 7—depth of the Moho surface (km); 8—state border; 9–10—directions of general and rotational movements of rocks. Letters on the maps: ABC—plates of the Eastern European Craton: A—Fenno-Scandia, B—Sarmatia, C—Volga-Uralia; tectonic structure: VM—Voronezh Massif, DDP—Dnieper-Donets Paleorift, US—Ukrainian shield. Segments of the Dnieper-Donets Paleorift: Ch—Chernihiv, Lh—Lokhvytsa, Iz—Izium; DF—Donets Foldbelt. Inter-block zones of the Ukrainian shield: I—Kropyvnitskiy; II—Ingulets-Kryvyi Rih-Kremenchug; III—Kamiaynsk; IV—Odessa; V—Central Priazov-Sloviyanogirsk. Suture belts and zones: IKK—Ingulets-Kryvyi Rih-Krupets`, AV—Alekseevka-Voronezh belts; Ya-Tr—Yadliv-Traktemyriv, K-Kr—Kryvyi Rih-Kremenchug, OP—Orikhiv-Pavlograd, by [2, 3, 12]. Frame of inversion deformations: 1—Kropyvnitskiy; 2—Ingulets-Kryvyi Rih; 3—Kamiaynsk; 4—Verkhovtsy-L’hov; 5—Kolomak-Kobeliaky; 6—Balakliya-Synelnykove; 7—Priazov-Slavyanogirsk, by [2, 13, 14, 15].

2. Stress and deformation sources of Earth’s crust of the Eastern European Craton

The stress field of the Earth is the result of a constant horizontal alignment of the maximum stress axis σ1 on continental lithosphere plates [17]. The established correlation between the direction of the stress axis and the vector of plate movement speed is evidence that the forces of plate interaction determine the stress field in intra-plate areas [17, 18]. The EEC has homogeneous collision compression due to its relatively homogeneous structure. The stress axis is oriented along the meridian, with the exception of Sarmatia, where it has a north-northwest orientation, associated with lithosphere anisotropy [12]. The stress field structure within EEC is determined by the sum of forces from various sources and depends on the degree of lithosphere anisotropy of constituent plates.

Contemporary and ancient stress fields are characterized by different levels of crust deformations. Contemporary field tensors are determined instrumentally [3]. Both strong and weak stresses, which do not cause residual deformation of rocks, are analyzed. The tensor of the ancient field is restored using the mechanical indicators of rocks and minerals, so only those stresses are determined whose magnitude exceeds the limit of irreversible deformations. The consequences of ancient deformations are observed only in the near-surface parts of the crust, but modern stresses can cover the entire upper consolidated crust with a thickness of 20–25 km. The deviator component of the field is taken into account—the excess stress, which is formed under the influence of certain forces in addition to litho-static (gravy) loading. If the stress field is stable for the upper part of the crust (20–25 km), the stress axes orientation on the horizontal with depth does not change [18]. The stress magnitude increases in consolidated rocks, so the stress deviator component is most often manifested in the basement and usually absent in the sedimentary cover of the platforms. The exception is the sheared folded mantles of intra-plate basins, where the stress field can be independent of the autochthonous platform they overlie [19].

Sources of stress causing deformation of the lithosphere are considered [15]: Planetary rotational forces, caused by changes in the parameters of the planet’s rotation and the Earth-Moon system; load forces on the Earth’s surface of the sedimentary thickness, tectonic and volcanic covers, negative and positive structures; local gravitational inhomogeneities, caused by variations in the density of mountain rock masses; forces arising during uneven bending of layers during heating/cooling of the rheological layered lithosphere; forces associated with the movements of lithosphere plates, including membrane forces, initiated by the movement of plates on the Geoid with variable curvature (“Plate driving forces”).

Strong stresses (up to 90 MPa) inside the plates can create mountain ridge loads [20]. However, the static pressure of excess masses creates anomalous fields only within the height of relief and does not affect the regular field of lithospheric plates. Stress inside the plates is influenced by movements at subduction borders: slab pull forces and trench suction [21]. However, the overall action of the plates is limited to the area adjacent to the subduction zone, so it does not significantly affect the upper part of the crust inside the lithospheric plates. The effect of drag and resistance forces that can act at the base of lithosphere plates (“Drag forces mantle drag”) is also insignificant [22]. A significant source of stress is the pressure of spreading oceanic ridges. Evidence of this is the intra-plate stress in the east of North America and in the northwest of Europe as a result of crustal divergence in the Mid-Atlantic Ridge [17, 18]. Regarding the influence of continental rifting, it has been established that against the background of crustal extension stages, episodes of compression and inversion uplift occur due to stress inversion in the lower layers of the lithosphere. Tectonic and sedimentary inversion of the rift is accompanied by changes in the thickness of the basin’s sedimentary cover without changing the overall movement of lithosphere plates [23]. Therefore, to model the evolution of the rift basin, the factors of dynamic interaction between the forces of the distant stress field from movements in collision belts and the loading/unloading of the basin structure through denudation and under the influence of static pressure of relief heights are taken into account. During the pauses between stress episodes up to 30 MPa, rift basins experience an uncompensated inversion of depths (up to 2 km) due to the cessation of debris inflow from the rising banks. This creates an additional force source for the formation of tectonic rifts during activation phases against the general divergence of continents.

Thus, Planetary and local forces act together, but the decisive contribution to the stressed state of the lithosphere is made by Planetary forces, the rest of the forces provide secondary variations. Collision stress propagates inside the plate for hundreds of km in EEC. The internal plate field is most influenced by induction stress from the movements of the Arabian and Scythian continental plates and the North Atlantic Ridge spreading. The Alpine-Himalayan collision belt, as a global structure-concentrator of stresses and deformations, has a greater influence on the state of the modern stress field in Earth’s crust of the EEC [7, 8, 9, 10].

3. Formation and evolution of the Sarmatia Earth’s crust

3.1 Anisotropy of lithosphere of the Sarmatia

The DDP is an element of the Sarmatia-Turan intraplate movable belt of the Eurasian Platform, extending from the Pripyat depression trough to the mountain-folded Tian-Shan [7]. The DDP crosses the Sarmatia obliquely, dividing the Archean-Proterozoic crystalline massifs of the USh and the VA (Figure 1) [2, 11, 12]. The structure of the Paleorift consists of an early Riphean Trog in the basement, a Devonian Graben and the overlying upper Paleozoic over-rift depression and Mesozoic-Cenozoic platform syncline (Figure 2) [13, 24]. The boundary faults of the Graben reach the Moho, form the tectonic shoulders, and control the early Riphean and Late Devonian rift complexes [24]. The Graben axis is convexly curved to the south and changes its stretch from sub-latitudinal to latitudinal in the southeast direction. In the same direction, the depths of the surface of the Archean-Proterozoic folded foundation increase from 5 to 20–22 km (Figures 1 and 3) [12, 13]. The Graben structure is diagonally superimposed on the pre-rift meridian seams (Kherson–Smolensk and Ingulets-Krivoy Rog, Donetsk–Bryansk) and crust-mantle suture (Odessa-Talniv, Ingulets-Bryansk, Verkhovtsy-L’gov, Odessa, and Central-Azov-Slovyansk) (Figure 1). The sub-meridian tectonic seams and sutures are fan-shaped and divided into component crust-mantle faults from the USh in the northern direction on the slopes of the VA (Figures 1 and 3) [25]. The transverse frame divides the structure from west to east into the Chernihiv, Lohvitsa, Izium, West Donets (segments), and the Donets Foldbelt (DF) known as the Paleo-basin “Donbas” (Figure 3). The Ingulets-Krivoy Rog-Krupets tectonic seam divides the VA into the Bryansk (West) and Kursk (East) micro-plates, and the USh into the West and East micro-plates. The Verkhovtsy-Lgov fault is the central lineament of the Ingulets-Krivoy Rog-Krupets seam, which breaks the Paleorift almost in half (Figures 1 and 3) [25]. Along the fault, the axis of tectonic symmetry formed, which separates the West and East micro-plates of Sarmatia of different structure, thickness and substance composition of the lithosphere layers, and opposite kinematics of displacements of rock [13]. The lithosphere under the Graben thins southeastward from 35 to 42 km in the Chernihiv to 15–25 km in the Lohvitsa and Izium segments, and thickens to 25–30 km at the border with the DF (Figure 4) [12]. The density of the crust at the Moho boundary increases southeastward from 3.12 g/cm³ in.

Figure 2.

Sections of the lithosphere East Europe Craton along the DOBRE fraction’99/DOBRE-2 profile, by [24]. Symbols: a—sedimentary cover, b—consolidated crust, c—mantle; 1—P-wave speed, 2—Vp/Vs, 3—reflection horizons, 4—Moho, 5—isotherm, grad C, 6—faults, 7, 8, 9—Earth’s crust: 7—upper (low and regular density), 8—middle, 9—bottom; 10—crust-mantle mix, 11—low and speed mantle, 12—deep fluids, 13—thermal anomalies, 15—roof of the asthenosphere: geothermal and seismic.

Figure 3.

Scheme of the Precambrian crystal basement relief of the Dnieper-Donets Paleorift, by [13]. Symbols in Figure 1.

Figure 4.

The deep structure of the Sarmatia Earth’s crust on sections across the Dnieper-Donets Paleorift, by [13]. Symbols: 1—the Moho boundary; 2–6—layers of the Earth’s crust: 2—sedimentary, 3—granite, 4—diorite, 5—basalt, 6—crust-mantle mixture; 7—edge faults of the Paleorift; 8—axis of the Paleorift.

Chernihiv to 3.18 g/cm³ in the Lohvitsa and from 3.18 to 3.20 g/cm³ in the Izium to 3.36 g/cm³ in the DF (Figure 5). The most powerful lithosphere up to 280 km lies under the Chernihiv, but under the Lohvitsa is a dome-like rift relic at the base of the lithosphere at a depth of 120 km. The base of the lithosphere rises from the Izium to DF to 80 km, which is due to the opening of the Devonian rift basin to the southeast in the Tethys Ocean [8, 9, 10].

Figure 5.

Scheme of the sole of the Dnieper-Donets Paleorift lithosphere and the density on the Moho boundary, by [13]. Symbols in Figure 1. Notations: The dashed line is the north boundary of the high-velocity horizontally in the Golitsyn-Heiko layer.

Due to structure-material differentiation in the mantle at depths of 50–250 km, zones of anomalous longitudinal seismic wave speeds were formed (Figure 4) [12]. In the zone of reduced speed under the Lohvitsa and the east of the Chernihiv, a “layering” of speed anomalies is localized. In the lithosphere under it, at depths of 70–130 km, a layer of anomalously high speed was formed through the intrusion of mantle matter from under the massif of the USh [12, 24]. The mantle subducted slab was facilitated by a shallow tilt of the northern side fault plane, which type changes with depth from a down-throw to a subduction-throw (Figure 2) [25]. In the west of the Chernihiv, the consolidated crust is made up of a “diorite layer,” and the subcrustal mantle has increased longitudinal seismic wave speeds (Figure 4). In the east, a typical “granite-free” rift crust is developed, and magnetic bodies are localized in the densified “diorite” and “basalt layers.” The relief of the Moho boundary coincides with the relief of the lithosphere base. In the Lohvitsa, the crust is “granite-free,” and the “basalt layer” also decreases in thickness due to the thickening of the “crust-mantle mixture” layer. In the roofs of the “basalt layer,” “crust-mantle mixture,” the bases of crust and seismic lithosphere, a dome-like rise is formed, which is a relic of the early Riphean rift.

The western part of the Izium has a “granite-free” crust with a thinned “diorite layer,” the “crust-mantle mixture” is ubiquitous (Figure 4). “Diorite” and “basalt” layers are densified and magnetic. Here a gradient rise of the Moho boundary and a shallow rise of the lithosphere base were formed. In the subcrustal mantle, a lens-shaped body with a density of 3.50 g/cm³ is localized (Figure 5). In the east part of the Izium, the crust contains a thin “diorite layer,” and a rise is traced in the roof of the “basalt layer.” The “crust-mantle mixture” forms a lens of abnormally high density (3.20 g/cm³) with a thickness of 12 km at the Moho boundary. Under the Moho boundary at depths of 60–80 km, the upper mantle is densified to 3.50 g/cm³. In the West Donets, the thinning of the “diorite layer” under the axial part of the Graben is associated with the maximum rise of the roof of the “basalt layer.” On the west slopes of the DF, sedimentary cover in the basement of the Graben lies on the “basalt layer.” The “crust-mantle mixture” layer reaches the maximum thickness of 16 km. On the Moho boundary in the Central-Azov-Slovyansk suture, a flexure of 5 km amplitude was formed (Figures 3–5).

Anisotropy of the Sarmatia lithosphere under the DDP is manifested from the surface to the base in variations in the composition and thickness of Earth’s crust layers, relief of the Moho, the lithosphere basement, and the consolidated crust (Figures 2–5). The longitudinal structural-material differentiation of the lithosphere is in the structure differences of the West and East micro-plates. An inhomogeneity in the orientation of the shear wave vector was established [12], which is considered evidence of structural differences, conditioned by the “built-in” (“frozen”) anisotropy of the lithosphere [25].

3.2 Geodynamic settings of the Sarmatia Earth’s crust deformations

The division of the EEC by intra-plate collision and movable belts into three lithosphere plates (Fenno-Scandia, Sarmatia, and Volga-Uralia) provides them with relative geodynamic independence and increases the overall and intra-platform tectonic mobility. At around 2.0 Ga subduction the oceanic plate, which separated Sarmatia from Volga-Uralia, has led to the collision with Fenno-scandia. The accretion of the two continental plates was completed at around 1.75 Ga. The Sarmatia is a stable Archean Craton, created by Late Archean-early Paleoproterozoic fusion of several different ages micro-plates ranging from 3.8 to 2.8 Ga [24]. Neoarchean-Proterozoic crust-mantle sutures intersect the 180–260 km thick Sarmatia lithosphere and sink into the mantle to 100–200 km (Figures 1–3). Sutures form a tectonic framework that controls the processes of Earth’s crust structural evolution due to repeated reactivation in a variable stress field. Within the USh, the DDP and the VA, displacements with horizontal amplitudes of tens of km occurred along this framework. In the Precambrian, the west and east parts of both crystalline massifs developed as separate micro-plates and were located far apart from each other. During the Vendian and Phanerozoic, the North pole moved from equatorial latitudes to the present position, and both micro-plates became part of Sarmatia [2, 3, 8, 9, 10] due to a 90° clockwise rotation of the EEC [3].

The model of the Precambrian evolution of Sarmatia covers several stages of collision with the subduction of continental crust and spreading with the divergence of the West and East micro-plates [3]. At Paleoproterozoic, collision of the Middle Dnieper and Azov Archean blocks were directed in the northeast and southwest of the current coordinates. Three continental (Rosin-Bug, Middle Dnieper, Azov) and two sub-oceanic micro-plates (Kryvyi Rih-Odessa, Dnieper-Azov) participated in the convergent process. Subduction of oceanic micro-plates under the Middle Dnieper and Azov micro-continents occurred with the closure of the Kryvyi Rih-Odessa paleo-basin. The movement of plates over the mantle plume ended with the form of the Holovaniv-Inhulets-Kryvyi Rih collisional suture. Due to the intrusion of the second plume, the Western and Eastern micro-plates diverged from the collision suture, forming a Trog between the Pervomaisk and West Inhulets collision zones. After the plume was extinguished, the second convergence of the West and East micro-plates of the Talne and Kryvyi Rih-Kremenchuk fault zones was formed. Under the influence of the third plume, the Holovaniv-Inhulets-Kryvyi Rih collision suture diverged, forming the Holovaniv and Inhulets-Kryvyi Rih sutures. The processes of crust convergence and divergence conditioned the anisotropy of the Sarmatia lithosphere through the formation of movable suture zones.

The model of tectonic and magmatic evolution of the DF includes several stages inherent to continental rifting [14]. The Riphean stage is associated with crustal stretching, the formation of a chain of Trogs, and the Sambets basaltic magma complex with gold-silver and copper mineralization. After the Early Paleozoic stabilization, rifting began in the Devonian with the reactivation of the old rift structure and the formation of a Graben belt. Volcano-plutonic magma rocks are composed of the Azov titan-pyroxenite and differentiated volcano complex from diamond-bearing alkali picrite to rhyolite. The synclinal stage in the Carboniferous involved submersion and the deposition of a coal-carbonate formation. Against the background of platform activation, the Anastasievka magmatic complex was formed, composed of gabbroid dikes and a volcano from basalt to alkaline rhyolite. After its completion, the Permian-Early Triassic stage of inversion uplift of the basin took place. It was accompanied by intrusions of granites, diorites of the shonkinite-monsonite-plagiophyr phase of the South Donbas complex. In the late Triassic, the stage of activation began, caused by movements and pressure from the south of the Caucasian micro-plates. It is associated with intrusions of andesites of the Nesvitaevka and lamprophyre Mius-Kerchik complexes. The stage of stabilization set in with the transition to platform mode at the boundary of the Cretaceous and Paleogene before the neo-tectonic activation. Fading tectonic movements were accompanied by diorite-dacitic dikes, low-temperature hydro-thermalites, and crushing of Lower Cretaceous rocks.

The kinematic model of rifting presupposes the formation according to the “athermal” mechanism through the elastic fracture of the “cold” continental Earth’s crust [14]. At the pre-rift collision stage in the Late Proterozoic, in the north-western diagonal faults at azimuths 291–312°, and 315–339°, a linear “germinal” strike-slip zone was formed. At the early Riphean in the shear stress field, movements along the zone initiated the formation of a chain of Trogs by the Pull-apart basin mechanism. At the Devonian rift stage, the plate was stretched, and the blocks of the Earth’s crust diverged south of the chain of trough depressions along dipping ledges, forming a Graben belt. Rifting initiated simple crust stretching deformations in the field with the horizontal orientation of the intermediate and minimum stress axes σ2, σ3 under the vertical axis of maximum stress σ1. Open to crust-mantle fluids, vertical detachment fractures, and faults formed in the rocks. Parallel to the extension of the rift belt was the axis of intermediate stress (σ2), orthogonal was the extension axis (σ3). Rifting was controlled by shifts-transforms in the northeast diagonal strike-slips in azimuth directions 24–30°, 39–45°, and 54–63°. Along with the transformations, “tectonic rift rails” were formed, along which the rift shoulders diverged in sub-latitudinal fault zones, and the segmentation of the future rift belt was laid. The tectonic position of the Devonian rift on the mobile outskirts of the EEC contributed to active inductive lithosphere deformations under the influence of the movements of the Tethys’s oceanic plates [8, 9, 10]. The rift’s southeast flank was located on the northern outskirts of Tethys and opened into the area of Late Paleozoic sub-oceanic basins and island arcs. During the Late Hercynian-Early Cimmerian activation, this area was transformed into the basement of the Scythian-Turan Plate. Platform paleo-basins known as “Dnieper-Donets depression” and “Donbas” existed from the Middle Devonian to the Early Permian [8, 9]. This is evidenced by relics of the rift structure, Devonian alkaline-basaltic volcano, and folded Paleozoic platform cover. Both basins ceased their existence due to fold deformation at the Late Hercynian Saal’ phase, about 0,250 Ga [24, 26].

The results of the study of stress fields of the Sarmatia [3, 8, 9, 10, 16, 26, 27, 28] indicate that tectonic inversion and folding in the southeastern segment of the DDP began at the end of the Early Perm (Saalian and Palatinate phases) of the Late Hercynian tectogenesis. It is believed that at the first stage, there were inversion deformations according to the model of oblique left-sided collision (Figure 6). They were influenced by the orogenic movements of the northern front of the collisional orogen formed on the active plate within the Paleotethys Ocean [8, 9]. In the body of the Sarmatian plate, this collision led to the formation of a stress field of tangential compression falling in the northeast direction.

Figure 6.

Schemes of geodynamic regimes of the Sarmatia at the tectonic inversion stage of the Dnieper-Donets Paleorift, by [29]: a—collision mode of the initial stage of inversion; b—local transtension in the mode of right shift of the main stage of inversion; in—the scheme of induction influence of the remote collision orogen. Symbols: 1—mobile area of the active environment; 2—deformation conditions: top—collisions, bottom—strike-slip deformations; 3—model trajectories of compression stresses; 4—modes: a—tangential compression; b—strike-slip deformation.

Mesozoic and Cenozoic tectonic movements within the Sarmatia plate are considered as a consequence of orogenic processes in the Black Sea-Caucasian segment of the Paleo-Ttethys folded orogen [8, 9, 10]. These processes caused the formation of a northwest tangential stress field in the Sarmatia Earth’s crust (Figure 6). At the neotectonic stage of geological evolution, the Paleorift structure continues to develop in geodynamic conditions of deep collisional compression and near-surface extension in the regional horizontal shear stress field [29, 30]. The mode of deep compression may have been caused by both the colliding geodynamic processes of the interaction of active plates and orogenic mobile belts on the northern edge of the Tethys paleo-ocean, as well as the intrusion at the late stage of riftogenesis of the mantle plume with uplift and structural and material processing of the base of the lithosphere [8, 9]. At the second stage of tectonic inversion, in the Mesozoic and Cenozoic the progression of the Donbas inversion uplift and its subsequent structural evolution as a DF took place in the regime of tangential compression caused by collisional stress from the Late Paleozoic Caucasus orogen [10, 27]. The progression of the general collision compression caused structural complications of Hercynian folding structure strike-slip parageneses of right-handed kinematics were formed on this territory.

Post-rift stages of the Sarmatia Earth’s crust evolution are characterized by interference fields and combined deformation modes. In the collision field with horizontal axes σ1, σ2, and vertical axis σ3, vertical shear fractures are impenetrable for fluid migration and laterally penetrable detachment fractures formed through over-thrusts. Break-down structures were formed in the basin’s basement by up-throws, “reverse” faults. Linear anticline folds were formed in the mantle, in orthogonal to the extension of which was the σ1 axis, and in parallel was the σ2 axis. Through the interference of collision stress modes and shifts in the field with horizontal axes σ1, σ3, and vertical axis σ2, the platform mantle underwent strike-slip and cover-thrust deformations. The geodynamic situation was characterized by a sub-meridian orientation of the stress axis with a variable inclination to the southwest (Saal’ phase), north (Laramian phase of the Early Alpine), and northeast (Attic phase of the Late Alpine) [16, 26]. The stress field was influenced by the pressure from the Donbas, which underwent tectonic uplift. In the Cimmerian (Donets phase) and Alpine (Laramian phase) activation, this one underwent cover folded deformations and transformed into the DF [8, 9, 10, 16, 26]. In the adjoined West Donets segment, the Mesozoic complex is eroded, three structural surfaces—Hercynian and two Alpine (Laramian and Attic), the structure of which is controlled by age-differentiated deformation grids of crust-mantle sutures [16, 26], were formed in the platform cover. Hercynian folding took place in the over-thrust-strike-slip mode of left-handed kinematics, while Cimmerian and Alpine cover-over-thrust in the strike-slip field of right-handed kinematics [8, 9, 10, 27]. In the Alpine flore of the West Donets segment and on the west slopes of the DF, under the overthrust framework, expressed in the Donets Ridge cover-throw scales were formed [16]. Displacement of blocks on the southern slope of the Graben caused the superimposition of sediment-volcano rift and platform cover at several stratigraphy levels by over-thrust plates of crystalline rocks. On the northern slope of the USh, the South Donbas Melange Zone was formed, manifested in relief as the Azov Heights [16].

Thus, the evolution of the Earth’s crust of the Sarmatia counts several significant stages due to the geodynamic situations and tectonic modes that arose on the EEC [2, 3, 6, 13, 16, 24, 28]:

1. Paleo-proterozoic collision of Fenno-Scandia and Sarmatia in the uplift mode with the unification of the crystalline massifs of the USh and the VA;

2. Early Riphean “athermal” rifting with the splitting of the “cold” continental crust along the linear germinal strike-slip zone, forming the chain of Trogs by the Pull-apart mechanism;

3. Devonian epi-continental “arch-thermal” rifting through the divergence of the crust under the influence of the mantle plume, forming the Graben sedimentary basin;

4. Hercynian-Early Cimmerian activation under collision conditions against the background of synchronous subsidence of the superimposed platform depression with the formation of linear folded zones involving salt tectonics;

5. Late Cimmerian-Alpine activation in the situation of interference of collisional stress and regional strike-slip field with territory uplift, basin desiccation and the formation of the West Donets Cover Folded Region and the DF.

3.3 Fault-block divisibility of Earth’s crust of the Sarmatia

In the Earth’s crust of the Sarmatia, eight azimuth directions of the Planetary fracture grid used by faults have been detected [2]. Faults of different directions in geochronology differ in genetic type, direction, and kinematics of displacement in geochronology: 285–290° (PR3, O, S—chips, strike-slips, ₽—over-thrusts, up-throws) 310–315°(D-C1—over-thrust, up-throws, N,—chips, strike-slips), 340–345° (PR3—gaps, throws, C1-2, C2-3, P2, T2-3—chips, strike-slips), 0° and 90° (D—chips, strike-slips) 15–20°(PR3, O, S—chips, strike-slips, ₽—gaps, throw), 45–50° (gaps, down-throw), 70–75° (PR3—over-thrust, up-throws, C1-2, C2-3, P2, T2-3—chips, strike-slips). Based on 3D seismic exploration data, the properties of tectonic lineaments of the DDP have been determined. Due to the statistical analysis of over 30,000 faults of the basement and four structural surfaces of the sedimentary cover, six realized azimuth directions of the Planetary grid have been established: 1—273–279° and 9–18°, 2—282–288° and 24–30°, 3—291–312° and 39–45°, 4—315–339° and 54–63°, 5—342–351° and 72–78°, 6—354–6° and 84–90° (Figure 7) [6, 15]. There are two directions (291–312°, 315–339°) in the northwest diagonal system of faults located along the extension of the Paleorift, creating a typical strike-slip pattern for the inner structure of the deformation zones (Figure 8). In the pre-rift stage, they laid the germinal linear shear zone.

Figure 7.

Pie charts of the spatio-temporal distribution of the stress field tensor of the Sarmatia Earth’s crust in the Dnipro-Donets Paleorift. A—scheme of the deformation ellipsoid tensor, by [3], B—the regional plan, C—the central part (Lokhvytsa, Izyium segments). The stress field tensor: б1—the maximum normal stresses; б2—the average normal stresses; б3—the minimum normal stresses; τ1, τ2—the paar of maximum tangential stresses.

Figure 8.

Comparative tectonic scheme of rift and post-rift fault frames and stress field tensor of the Sarmatia Earth’s crust in the Dnieper-Donets Paleorift in the Precambrian basement, sedimentary cover, and topography. Insets: a, d—stress field tensor: a—rift stage; d—collision stage (Attica); b, c, e—diagrams of the azimuthal distribution of geodynamic axes: b—in the foundation; c—sedimentary cover, e—in relief.

There are three directions (24–30°, 39–45°, and 54–63°) in the northeast diagonal system formed shifts-transforms, which controlled the process of rifting. The faults of the diagonal and orthogonal system in azimuth directions (273–279°, 291–312°,315–339°, 354–6°, 9–18°, 39–45°, 54–63°, and 84–90°) changed structural-dynamic manifestation at the stages of platform activation (Figure 9) [6, 13]. The “reverse systems” faults were part of the tectonic grids of folded layers of different ages in the basin cover. Changes in the dynamics and kinematics of “inverse systems” occurred in a new geodynamic setting due to space–time inversions of the stress field tensor. Inversion deformations initiated changes in the tectonic mode and rebuilt the rift structure with a new plan controlled by a younger tectonic framework.

Figure 9.

Periodically scheme of the kinematic mechanism of the spatio-temporal inversion of the stress field tensor of the Sarmatia and the formation of the deformations structural plans of the Dnieper-Donets Paleorift in Phanerozoic, by [6]. Symbol: R—“reverse faults”; another symbol in Figure 7.

Thus, there are three structure plans of deformation at the Phanerozoic stages of structural evolution of the Sarmatia Earth’s crust were formed: the rift (D2-C1t), the platform syncline (C1v – P1), and the collision (P2-T—Q) (Figure 9) [6, 13, 14].

3.4 Distribution of inversion deformations in the Dnieper-Donets Paleorift

It is a priori assumed tectonic stresses and deformations during the platform activation first affected the Precambrian basement of the Sarmatia as a base layer of the Earth’s crust [2, 6, 13, 14]. Later, the deformations covered the basin sedimentary cover [6, 16, 26, 29]. The deformation of the rift structure increases along to the southeast. Rift throw relics remained in the West micro-plate, and on the northern flank of the East micro-plate along sharp changes in basement depth, reflected by elongated gravity field gradient anomalies. The structural boundary between weakly dislocated (Chernihiv, Lokhvytsia, Izium) and completely inverted (West Donets, DF) segments is drawn along the Central-Azov-Slovyansk crust-mantle suture at the meridian of Balaklia city (Figure 10) [29]. Here, the north (Kochubiyevka-Volvenkovo) and south (Sosnivka-Stepkovka) chains of axial salt dome anticlines are interrupted (Figure 11). To the east of the border extends the only axial Druzhkivka-Komyshuvakha salt wall. The rift framework and Hercynian folded lines were segmented by diagonal strike-slips into coulisse sections. Border throws bend in the plan initially to the east, and at the meridian of Donetsk city to the southeast, as a result of which the area of the Graben expands in the West Donets segment. Due to Hercynian movements, the rift structure underwent an uplift of several km and linear folding. In the north of the West Donets, the vector of Hercynian movements ascends to the southwest and crosses at a large angle the eroded surface of the basement. This contributed to the over-thrusting of Precambrian basement rocks onto the Paleozoic cover along the gentle North Donets overthrust (Figures 12 and 13). In the south Graben border with the Azov Massif of the USh, the basement surface serves as a plane of tectonic ruptures due to a northeast azimuth of fall coinciding with the Hercynian movement direction (Figures 6 and 14). By the same kinematic mechanism, the South Donbas Melange Zone in the Paleozoic complex has been formed. The Melange Zone did not develop further west due to a structural barrier—a ledge of basement under the southern slope of the Graben (Figures 13–15). The west slope of the ledge is bypassed by an ensemble of small horst and grabens of rift stretch. Their axes turn from south to southeast, rise up, and cut at sharp angles at tectonic contact with the Zone of Melange. The rift stretch structure is orthogonally superimposed by the frame of unidirectional “unconformity” down-throws. Because of this, the hanging blocks at the base of the grabens lie 2–3 km higher than the foot blocks. The rise in the pre-Triassic erosion level to the southwest indicates pre-Mesozoic denudation from the southern slope of the Graben to the Azov massif with a vertical amplitude of more than 3 km.

Figure 10.

The map of the pre-rift faults of the Dnieper-Donets Paleorift Precambrian basement, by [31]. Symbols: 1—the sedimentary cover local structures; 2—salt doms; 3–5—fields: 3—gas; 4—gas and condensate; 5—oil; 6–11—faults: 6—of the borders, 7—of the basement, of the Devonian, of the Carboniferous, of the Permian, of the local structures; 12–13—the changes of the sedimentary complexes thickness: 12—by stratigraphic distribution, 13—by vertical amplitudes and dip; 14—boundary zones of fold deformations; 15—boundaries of the West-Donets Cover Folded Region.

Figure 11.

Fragment of the tectonic map of the Dnieper-Donets Basin, by [31]. Symbols: 1—the Precambrian basement surface marks, m; 2—seismic profiles: a—old, b—new, c—reinterpreted; 3—deep wells: #/mark; 4—faults: a—throw, b—thrusts, c–d—diagonal strike-slips of extension: c—north-west, d—north-east; 5—boundaries of the West-Donets Cover Folded Region; 6—anti- and sin-structures; 7—Devonian salt domes; 8—boundaries of the tectonic zones; 9—the basin boundaries; 10—compensatory throughs, 11—depressions, 12—anticlines, 13—anticlines shaft, 14—salt-dome anticlines zones: 1—Kochubiyvka-Volvenkovo, 2—Sosnivka-Stepkivka, 3—Druzhkivka-Komishuvakha, 15—ledges.

Figure 12.

Regional geologic sections of the Dnieper-Donets Basin, by [30]. A1A2—through the south-eastern part of the Izium segment of the Dnieper-Donets Basin. B1B2—through the central part of the West Donets segment.

Figure 13.

Pokrovs’k-Biryukovo seismic section through the south slope of the West Donets Graben and the axial part of Donets Foldbelt, by [29]. The inset is the profile line on the map (Figure 11).

Figure 14.

Scheme of the West-Donets Cover Folded Region tectonic frame. Symbols: 1—north border of the region; 2–4—thrusts: 2—Herzinian, 3—Laramian; 4—Attic; 5—low-dislocated Paleozoic autochthon; 6—Herzinian neo-autochthon; 7—the Priazov Massif; 8—the South Donbas Melange Zone; 9—West-Donets Cover Folded Region: (1) Luhansk-Komyshuvakha thrust folding area, (2) Kalmius-Torets Scally of Thrust Area; 10—the Donets Foldbelt. Insert—scheme of the cover folded system.

Figure 15.

Kinematic model of the Dnieper-Donets Paleorift inversion deformation of Earth’s crust on the map of vertical neo-tectonic (Alpine, Holocene) movements. Symbols: 1—pre-rift deep faults; 2—ring anomalies on satellite images; 3—neo-tectonics strike-slip zones, 4—the Verkhovtsy-Lgov tectonic axis of kinematic symmetry; 5–6—structural parageneses of deformations: 5—tension, 6—compression; 7–8—rocks displacement directions; 9—the frontal part of the Nizhyn-Ichnia structural paragenesis of longitudinal bending; 10—the most dislocated area of the basin; 11—West-Donets Cover Folded Region; 12—the frontal part of the indenter horst mega-block of the Donetsk Foldbelt. Another symbol in Figure 1.

The West Donets Cover Folded Region was formed in the Hercynian and Alpine folded floors in the West-Donets segment and on the west slopes of the DF (Figure 11). The wedge-shaped region between Luhansk, Izium, and Donetsk cities consists of echelons of scales of tectonic covers and coulisse-jointed anticlines. The multi-age strike-slip and thrust frame controls the “cross-thrust” structure of deformations (Figures 12 and 14). The north flank of the Region is bounded by the North-Donets thrust, in the south the Samara and Novoselyvka thrusts separate it from the Mélange zone. The Region is divided in half into two structural areas: (1) Luhansk-Komyshuvakha Thrust Folding Area, (2) Kalmius-Torets Thrust Scally Area by the axial Sulino-Kostyantynivka fault. In the axial zone of the Graben, echelons of over-thrusts control tectonic blocks of elongated east and northeast stretch, complicating the structure of the Bakhmut, Komyshuvakha-Liman, and Kalmius-Torets depressions. In the front of the blocks, thrust anticlines were formed with steep northeast wings and gentle southwest ones, which turn into monoclines in the rear. On the west flank, the sole of the Mesozoic cover plunges to the west to a depth of over 3 km within the Orchik depression in the Izium segment.

The Cover Folded Region is located above a triangular depression in the basement, filled with a sedimentary thickness of up to 20–22 km (Figure 10). In the center of the Region, in the zone of the Axial (Sulin-Kostyantynivka) upthrow-shear, the largest Main (Horlivka) anticline is located, with a length of 170 km at a width of 10 km (Figures 11–14). The main anticline with a ridge hinge was formed in the mode of lateral over-compression through the displacement of wings along the left-lateral up-throw and strike-slip with an amplitude of tens of kilometers. In the southeast direction, the hinge of the fold gradually rises by 8–9 km, as a result of which Mesozoic deposits on the wings are eroded. On the northern flank of the orocline, in the hanging wings of Hercynian thrusts (Drobyshivka, North Donets, and Matross), coulisse-sectioned chains of upthrow-anticlines were formed (Figures 11, 12, and 15).

The Late Alpine (Oligocene-Miocene) deformations are controlled by an over-thrust framework, behind which large blocks of the pre-Cenozoic complex are segmented into plates from the first to dozens of km and over-thrusting in the northeast direction (Figures 12 and 15). The maximum amplitude of horizontal displacements of rock blocks reaches 4.5 km with a vertical one up to 1.5 km. Attic deformations are only absent in the southwest part of the West Donets segment, they are significant in the rest of the territory. To the north and east, the Cenozoic surface rises sharply over 250 m toward the Donets Ridge. Altitude differences exceed hundreds of meters, corresponding to the amplitudes of displacements along the Attic thrusts. Attic movements are manifested in right-lateral movements along the Samara, Novoselivka over-thrusts in the south and the North Donets one in the north, and the formation of the Donets Ridge and the Azov Highlands.

3.5 Kinematic mechanisms of deformation of the Earth’s crust

According to instrumental determinations in the DDP and its outskirts, basement faults are mostly up-throw and strike-slips [3, 16, 26, 27, 28, 31]. In the sedimentary cover, in the shear field above them, rows of diagonal strike-slips were formed in correspondence to the coulisse arrangement of fracturing chips [5, 13, 29]. Due to the combination of elementary deformation modes (compression/expansion/displacement) in the fault zones, parageneses of “flower” and “duplex” structures are formed [32]. In the trans-tension mode, “tulip” and “contractional duplex” structures are formed in paragenesis of upthrow, strike-slipes, and overthrusts. In the extension mode, “palm tree” and “extensional duplex” structures are formed in paragenesis of throws, strike-slipes. Lateral displacements of rocks are controlled by synthetic and antithetic strike-slips, which are composed of R and R’-fracturing [7, 32]. Structure dislocation is formed along shears both co- and counter to the movement direction in the wings of faults and due to their combination, with a rotation and without it [7].

During the Hercynian collision, the rock displacement within the rift basin occurred toward the “geodynamic shadow” zones [29]. Evidence of longitudinal extrusion of rocks is the stress measurements on volcanic structures in the articulation zone of the FD with the Azov Heights [28]. In trachytes, two fields of shear type were found: Late Alpine (Attic) with a sublatitudinal compression axis and a submeridional extension axis, and early Alpine (Laramian) with the reverse location of the stress axes. In andesites, diagonal zones of slating of northwest and northeast extensions are established. In separate bodies, a late Hercynian strike-slip field of right kinematics is formed. A geophysical feature of movements along faults is stock-like zones of loss of correlation of seismic reflective horizons in the wave field due to tectonic fragmentation and loosening of rocks in the fault zones [5]. Evidence is a large stock structure in the seismic field at the hinge of the main anticline (Figure 13).

Hercynian movements caused the displacement of the masses of rocks from over-compressed areas in the central part of the rift to the largest by area and less compressed West Donets segment. The discharge area of the extruded masses widens in a wedge-shaped manner along the course of the tectonic flow from the extrusion area at the eastern border of the Izium segment to the western slopes of the Donbas (Figures 11 and 14). The orocline outlines of the unloaded massif coincide with a triangular depression in the basement filled with 20–22 km of rocks. Above the depression, there is an ensemble of arcuate bent, coulisse-linked, and strongly compressed upthrow-anticlines, transverse in extension to the direction of flow. Transverse folding is diagonally superimposed on the lines of large Hercynian anticlines of rift orientation (Figure 11). The apparent reverse overthrust of upthrow-anticlines backward along the flow indicates a lack of geological space at the point of its origin. Due to the situation of volume compression, the upper layers of allochthon lagged behind the general movement of the flow, therefore a structure of underthrust under transverse folds was formed here, not vice versa. Strong stress in this area initiated a high orogenic uplift of layers, the formation of secondary echelon overthrust-folding, and at least three intersecting structural levels (Hercynian and two Alpine). The masses extruded with the flow accumulated here synchronously with loading. Evidence of this is the intersection of the fault-thrust frame and zones of near-fault folding in the west—at the starting point of the flow origin (Figure 14). Further, along the flow, the excess of masses gradually dissipated due to stress reduction. Therefore, transverse thrust structures arise, and longitudinal thrust-folds are replaced by transverse ones. As the orocline expands, large folds elongate and turn into anticlinal swells. This is evidence of their formation due to tectonic disruption and thrusting onto a paleo-depression on the western slope of the DF.

Alpine movements were directed northwest toward the central part of the inverted belt. The lateral stress was induced by the movements of the mountain-folded Caucasus and the inversion uplift of the Donbas, which turned into a folded belt. Alpinotype deformations occurred due to the movement of masses from the west slopes of the DF along the previously formed orocline framework (Figure 14). Under the pressure from the east of the megablock-indenter of Donbas, the masses flowed in the opposite direction to the Hercynian current (Figure 15). As a result of their accumulation on the Hercynian neo-autochthon, two alpine (Laramian and Attic) systems of thrust covers were formed (Figures 12–14). This situation led to a significant thickening of the sedimentary cover in the literally limited space of the paleo-basin due to the thrusting of younger allochthon plates (Figures 14 and 15). The allochthon composed a post-sedimentation part of folded surfaces of the cover that was not compensated by sedimentation. According to the kinematic mechanism of transverse protrusion along dynamically conjugated and kinematic antithetic diagonal up-throw-slips without a rotational component of movements, a “cross-over thrust” structure of the “orocline of transverse protrusion” was formed (Figure 16) [7, 29]. The Cover Folded Region and the Melange Zone were formed on the south slope of the Donbas. The Melange Zone is expressed in the relief as the Azov Highlands stretches diagonally to the Donets Ridge and has similar elevation marks.

Figure 16.

Kinematic model of formation of the West Donets Cover Folded Region. Symbols: 1—advanced compression fan of the tectonic oroclin; 2–4—thrusts: 2—Hercynian, 3—Laramian; 4—Attic; 5—rift relics throw; 6—rocks displacement directions; 7—South Donbas Melange Zone; 8—West-Donets Cover Folded Region: (1) Luhansk-Komyshuvakha Thrust Folding Area, (2) Kalmius-Torets Thrust Scally Area; 9—anticlines: 10—sinclines; 11—linear strike-slip zones: [1]—north-east flank, [2]—Axial, [3]—south-west flank and kinematic of displacements. Inset: a—scheme of the “Longitudinal Extrusion Orocline”; b—scheme of the cover folded system.

The sublatitudinal rock displacement along the fault frame caused deformations of the longitudinal elongation of the rift structure due to the stretching of the rock layers along their strike [4, 7]. According to the first kinematic mechanism, the structure is divided into several tectonic blocks by rows of antithetical shifts of the same kinematics. The blocks jointly rotate like a row of dominoes, along shifts diagonal to the axis of the structure (Figure 17). Shear displacements are directed against the movement of the general flow of rock masses. The ensemble of dominoe blocks experiences unidirectional movements orthogonal to the stretching of their curved wings, which leads to an elongation of the original structure along its length [7]. In the West micro-plate on the southern slope of the Graben, a behind-the-scenes ensemble was identified, consisting of the Chernigov, Nezhin-Ichnya, and Priluky-Lubny dynamically related structures of longitudinal extension (Figure 18).

Figure 17.

The internal kinematics of the frontal part of the Nizhyn-Ichnia parageneses of the longitudinal bending in the central part of the Paleorift on the map of Holocene vertical movements. The block “dominoes structure” was formed due to the longitudinal displacement with the rotation of the blocks and on diagonal shifts of the same kinematics. Inset: A kinematic scheme of a block’s “dominoes structure,” by [7]; A, B—before and after deformation; C—a structural pattern.

Figure 18.

Model of a longitudinal bend kinematic mechanism of rift structure deformation in the south part of the Paleorift on the map of Holocene vertical movements. A number of the oroclines of longitudinal bending: ①—Chernihiv; ②—Nizhyn-Ichnia; ③—Pryluky-Lubny. Inset: Kinematic scheme, by [7]. Symbols (Figures 17 and 18): 1—basement faults; 2—basement surface marks, m; 3—strike-slip-throw frame of the oroclines; 4—edge shear-fault; 5—shear-transforms; 6—rocks displacement directions; 7—directions of extrusion and shear flow; 8—vectors of local tension; 9—arcuate fault-spreading; 10—the Verkhovtsy-L’gov deep fault; 11—tectonic axis of kinematic symmetry; 12—wells.

According to another mechanism, longitudinal tension is realized through multidirectional motions of blocks with rotation along two dynamically coupled diagonal strike-slips [7]. Due to the tectonic spreading of the rocks in opposite directions, a new intermediate trapezoidal block is formed in the fault zone (Figure 19). Subsequently, the rocks are squeezed out of the block, which experiences tensile deformation. Tension zones are formed in the basement, depressions are formed in the sedimentary cover above them. Such a geodynamic situation developed in the central part of the Graben at the boundary of the West and East micro-plates in the zone of dynamic influence of the Verkhovtsy-Lgov crustal-mantle fault (Figure 15). As a result of displacements of rocks in its wings, diagonal strike-slips like as a plumage were formed. Due to multidirectional movements with rotation along diagonal strike-slips, intermediate tectonic blocks were formed. The situation of divergence of the Earth’s crust initiated the spreading of crystalline rocks in opposite directions outside the blocks. Tensile deformations initiated the formation of an ensemble of local lenticular depressions in the basement surface, elongated along the meridian. Above them, in the sedimentary cover, the Vorskla depression was formed by the Pull-apart mechanism (Figure 20) [30]. The geodynamic region of the Earth’s crust extension is controlled by the strike-slip paragenesis of the tectonic “longitudinal extension fan” turned to the north (Figure 21) [33].

Figure 19.

Model of a “Transverse Spreading Fan” kinematic mechanism of rift structure deformation in the central part of the Paleorift on the map of Holocene vertical movements. Inset: A, B, C—before and after deformation; D—a structural pattern. Number of the structural paragenesises: ①—the Mirhorod orocline of longitudinal bending; ②—the Pereshchepino orocline of longitudinal bending; ③—the Vorskla lenticular pull-apart basin; ④—the Verkhovtsy-L’gov deep fault. Symbols in Figure 18.

Figure 20.

Model of the “transverse spreading fan” kinematic mechanism on a 3D tectonic reconstruction at the end of the Late Hercynian Saal’ phase in a section through the Vorskla depression in the central part of the Paleorift.

Figure 21.

Kinematic scheme of inversion complicated of Dnieper-Donets Paleorift Earth’s crust: 1—regional linear shear zones; 2—the Verkhovtsy-L’gov tectonic axis of kinematic symmetry; 3–4—folded longitudinal zones: 3—arcuate fault-spreading, 4—strike-slip-throw frame; 5–6—direction of blocks displacement and rotation.

At the collision mode, synchronous deformations of lateral tension and vertical extension also occurred [7]. In the basin confined space, under the influence of lateral stress, there was a tectonic extrusion of rocks upward free from compression of the ground surface. The uplift of masses on the surface initiated the formation of stress zones and uplifts over the fault zones as a source of erosion and wear in the basin. Examples are the Olenivka, Volodarka, Uspenka, and Rovenets over-thrust anticlines, which are reflected in the heights of the West Donbas relief (Figure 13). At the same time, a compensatory geodynamic factor caused the formation of extension zones of the Earth’s crust, dynamically and spatially conjugated with stress zones [2, 3, 15, 27]. As a result, “overlaid” depressions were formed, where the products of erosion accumulated, removed from nearby hills. Examples are from the largest Kalmius-Torets, Bakhmut, to the smallest Mykhailivka, Kuybyshevka, and other depressions are reflected in the basin’s relief (Figures 12 and 13). Modern uplifts and associated troughs are located diagonally to the strike of basement fault zones due to horizontal displacements of rock masses in the limbs of strike-slips in the platform cover (Figure 11). Together they form a segmented paragenesis of longitudinal elongation.

Thus, at the inversion stage, the rift basin underwent uplift and folding deformations, transforming the rift and post-rift platform complexes into folded structural floors of the cover. The tectonic framework of deformations formed by remobilized pre-rift crustal-mantle faults is superimposed diagonally and orthogonally on the network of rift throws. The strike-slip and uplift network controls the regional inversion structures of “longitudinal extrusion orocline of rocks” in the southeastern part and “the transverse extension fan” in the central part of the DDP. Therefore, paleo-basins “Dnieper-Donetsk Depression” and “Donbass” of the Paleozoic complex at the inversion stage of the evolution transformed into the West Donetsk Cover Folded Region and DF due to Hercynian and Early Alpine deformations. Currently, the inversion structure of the Paleorift is overlain by Mesozoic and Cenozoic complexes of the platform cover and additionally complicated by Alpine dislocations.

4. Conclusions

1. The stress field of Sarmatia was influenced by inductive movements of the Arabia and Scythia, spreading of the North Atlantic Ridge, pressure from the Caucasus and the indentor mega-block of the FD. Anisotropy of the Sarmatia is manifested from surface to base of the lithosphere in variations of composition and thickness of layers of consolidated crust, the Moho, the lithosphere, and the crust base. The structural-material differentiation of the lithosphere determined the processes of evolution of the Sarmatia Earth’s crust, including the division into the West and East micro-plates, the formation of a rift structure, and the inversion deformations of the DDP. A defining factor of the evolution of Earth’s crust is the remobilization of faults during platform activation and the formation of a deformation framework.

2. The deformation framework was formed along the pre-rift crust-mantle faults, diagonally superimposed on the rift fault structure frame. There are three plans of deformation were formed: rift (D2–C1t), platform syneclise (C1v–P1), and collision (P2–Q) in the Paleorift. The intensity of deformations increases along the rift strike to the southeast. So, in the West segments relics of the rift are preserved, but in the East ones, the structure is destroyed by Alpine deformations. Deformations are manifested in the formation of tectonic detachments, upthrows in the basement; coulisse strike-slip deformation framework of the sedimentary cover; “cross-overlapping over-thrust” structure of cover folded deformations of the Hercynian and Alpine; thickening of the cover in the limited space of the basin due to the over-thrusting of younger tectonic allochthons that make up the uncompensated by sedimentation part of the cover. The regional collision framework of deformations is formed by the parageneses “orocline of longitudinal extrusion of rocks” in the southeastern part and “fan of transverse spreading of rocks” in the central part of the Paleorift.

3. The paleo-basins “Dnieper-Donets Depression” and “Donbas” by Hercynian and Cimmerian deformations of the Paleozoic complex turned into the West Donets Cover Folded Region and Donets Foldbelt, were buried under the Mesozoic and Cenozoic platform cover and were further complicated through by Alpine deformations. Considering the ubiquitous distribution of inversion deformations, a conclusion is made about the transformation of the DDP into a Dnieper-Donetsk Movable Foldbelt. Neo-tectonic deformations are structurally manifested by the Donetsk Ridge and the Azov Highlands. Considering the ubiquitous distribution of inversion deformations, a conclusion is the transformation of the DDP into a Dnieper-Donets Movable Foldbelt.

Acknowledgments

I thank V. Karazin Kharkiv National University, Ukrainian Research Institute of Natural Gasses, Geological Enterprises of Ukraine for providing materials. The author is grateful to S.V. Goriaynov, V.I. Alekhin, V.A. Korchemagin, O.B. Gintov for sharing their measurements of stress fields in Ukraine, S.V. Goryainov for providing his sections and the structural map of the West Donbas Hercinides, V.I. Starostenko for the schemes of the Sarmatia lithosphere, Polivtsev A.V. for the map of vertical neo-tectonic movements.