Structural development and stress evolution of an arcuate fold-and-thrust system, southwestern Greater Caucasus, Republic of Georgia

Journal of Asian Earth Sciences 156 (2018) 226–245

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Structural development and stress evolution of an arcuate fold-and-thrust system, southwestern Greater Caucasus, Republic of Georgia

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A. Tibaldia, , F.L. Bonalia, E. Russoa, F.A. Pasquarè Mariottob a b

Department of Earth and Environmental Sciences, University of Milan Bicocca, Milan, Italy Department of Theoretical and Applied Sciences, University of Insubria, Varese, Italy

A R T I C L E I N F O

A B S T R A C T

Keywords: Greater Caucasus Faults Paleostress Folds

The southern front of the Greater Caucasus is quite rectilinear in plan view, with the exception of part of the Rioni Basin, where marine and continental deposits of Cretaceous-Neogene age were locally folded and uplifted; this resulted in the formation of an arcuate fold-and-thrust system that extends 45 km into the foreland. Although previous studies suggested that this system has developed only since Miocene times, our new detailed and systematic field measurements of brittle and ductile structures show a very complex history, consisting in four main phases of brittle deformation and folding, dated from Eocene to Quaternary times. We collected microtectonic data at 248 faults, and calculated the related paleostress tensors. The first two phases which we document here, predated folding and were characterised by dominant transcurrent faulting and subordinate reverse motions; the greatest principal stress σ1 was perpendicular and later parallel to the mountain belt. Afterwards, NW-SE, E-W and NE-SW trending, south-vergent asymmetrical folds started to form. In the western sector of the study area, folds are sinuous in plan view, whereas to the east they show a left-stepping, en-échelon geometry. Another two, brittle deformation phases took place after the folding, due to the activity of a set of right-lateral, strike-slip faults that strike NW-SE and NE-SW, respectively, as well as by left-lateral strike-slip faults, mostly striking NW-SE, NE-SW and NNE-SSW. These two additional phases were produced by a NE-SW to N-S trending σ1. The arcuate belt is marked by along-strike variations in the tectonic regime and deformation geometry, plus belt-parallel stretching. Based on our field data, integrated with published analogue models, we suggest a possible explanation for the Rioni structure, in terms of the oblique, asymmetric indentation of an upper crustal blocks moving to the SSW.

1. Introduction The Greater Caucasus is an orogenic system that formed in Cenozoic times following the Arabian-Eurasian plate collision. The orogeny developed owing to the formation of mainly WNW-ESE trending, folds and faults, which resulted in a rather rectilinear mountain front in plan view, especially along its southern side (Fig. 1). The Transcaucasian intermontane valley, which separates the Greater Caucasus to the north from the Lesser Caucasus to the south, is made up of: (i) the Rioni foreland fold-and-thrust belt and the Rioni foreland basin to the west (Adamia et al., 2011b; Forte et al., 2014), (ii) the Dzirula massif in the centre (Khain, 1975), and (iii) the Kura Basin to the east. The Kura and Rioni basins developed in Oligocene-early Miocene times, and they were then partially involved into the orogenic fold-and-thrust belts (Adamia et al., 1977, 2010; Banks et al., 1997; Mosar et al., 2010; Sosson et al., 2010a,b, 2013; Forte et al., 2010; Alania et al., 2016). The Rioni fold-and-thrust belt is made up of a complex pattern of structures

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that result in an arcuate zone of deformation extending 45 km to the SSW from the Greater Caucasus foothills (Fig. 1). Some of these structures are still active, as documented by seismological data (Tsereteli et al., 2016) and paleoseismological research (Tibaldi et al., 2017a, 2017b). In spite of the complex geometry of this sector of the Greater Caucasus frontal system, as well as the occurrence of active deformation that may be associated with potential seismic hazard, this area has received very little attention so far. In the international literature, a preliminary paper describing the general structure of the area was published by Philip et al. (1989), who documented the distribution of the main folds and presented paleostress data gathered at seven sites. Later on, Tibaldi et al. (2017a) performed new measurements along striated faults, and related paleostress inversion, at nine new sites. In regard to geophysical data, Banks et al. (1997) presented a geologicalstructural cross-section based on seismic exploration data, running across the Rioni Basin; on the other hand, Tibaldi et al. (2017b)

Corresponding author at: University of Milan Bicocca, P. della Scienza 4, 20126 Milan, Italy. E-mail address: [email protected] (A. Tibaldi).

https://doi.org/10.1016/j.jseaes.2018.01.025 Received 9 August 2017; Received in revised form 19 January 2018; Accepted 24 January 2018 Available online 06 February 2018 1367-9120/ © 2018 Elsevier Ltd. All rights reserved.

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Fig. 1. Tectonic map of the Arabia-Eurasia collision zone (modified after Sosson et al., 2010a, 2016). Abbreviations: AT Achara-Trialeti; LC Lesser Caucasus; EAF Eastern Anatolian Fault; NAF North Anatolian Fault; IAES Izmir-Ankara-Erzincan Suture.

trending greatest principal stress during the older phases and by a N-S to NNE-SSW trending greatest principal stress during the younger phase, as suggested by Tibaldi et al. (2017a), (ii) is compression parallel to the orogen reflected in the folding processes, (iii) how can the rotation of the stress axes be explained, (iv) how did the curvature of the fold-and-thrust system develop, (v) can field-based, microtectonic data confirm the active stress state as derived by earthquake focal mechanisms and thus contribute to seismic hazard assessment? The hereby provided data may serve to enhance understanding of how arcuate mountain belts develop, to gain a better knowledge of one of the main orogenic system of Earth, which is still poorly studied, and to improve the database useful for seismic hazard evaluation in a broad area that hosts several towns as well as the Enguri Hydroelectrical Scheme, the main energy production facility in the Republic of Georgia.

provided a number of seismic reflection sections across a main anticline known as Tsaishi fold, and Tsereteli et al. (2016) analysed earthquake distribution and focal mechanism solutions. The development of arcuate fold-and-thrust belts has been a major focus of study since the early 90s (Suess, 1909; Hobbs, 1914); this kind of characteristic pattern was referred to as “salient” by Miser (1932), because the structures are convex in plan view and extend farther towards the foreland. Many studies performed at salients suggest that materials are transported in three dimensions (Sussman et al., 2004), so that accurate analyses need to be carried out around and across the arcuate belts, in order to correctly reconstruct their overall architecture and strain pattern. Recent papers have contributed to elucidating the development of salients by way of kinematic field data (e.g. Platt et al., 1989), paleomagnetic data (e.g. Sussman et al., 2004), and analogue modelling (e.g. Lickorish et al., 2002; Calignano et al., 2017, and references therein). In some cases, kinematic data from fault-slip analysis show a quite simple and consistent picture of the development of arcuate belts, such as is the case of the Pannonian Belt. Here, the stress and strain directions of the arcuate fold-and-thrust systems allow to document an overall NW–SE compression, with fanning of the subhorizontal contraction axes from E–W to NNW–SSE (Zweigel et al., 1998). On the other hand, at the Rioni Basin, preliminary fault-slip data indicate a much more complex setting (Tibaldi et al., 2017a), with a major rotation of the horizontal compression axis and the widespread presence of strike-slip faults. All the above indicates that the Rioni Basin, at the foothill of the Greater Caucasus, represents an outstanding example of a recent salient, marked by a complex deformation history in spite of its short (Miocene-Quaternary) history. In view of the above, we carried out a detailed field survey of the whole Rioni area affected by compression tectonics, in order to unravel the various phases of deformation that affected it. Regional observations were integrated with systematic collection of microtectonic data, aimed at reconstructing the geometry of the folds and the evolution of stress orientations. In doing so, we wish to contribute to addressing a number of challenging scientific questions, including the following: (i) can the regional data confirm the complex evolution of the local stress field, characterised by a NW-SE and NE-SW

2. Geological background 2.1. Regional geology During the Neotethys subduction, several domains formed in backarc locations within the Eurasian Plate, among which the Greater Caucasus basin that developed in Early–Middle Jurassic times (Khain, 1975; Dercourt et al., 1986; Adamia et al., 1981, 2011a; Sosson et al., 2016), followed by the opening of the western and eastern Black Sea basins during the Cretaceous and Cenozoic (Adamia et al., 1981; Letouzey et al., 1977; Finetti et al., 1988; Zonenshain and Le Pichon, 1986; Okay et al., 1994, 2013; Robinson et al., 1996; Spadini et al., 1996; Cloetingh et al., 2003; Vincent et al., 2005; Yegorova and Gobarenko, 2010; Khriachtchevskaia et al., 2010; Stephenson and Schellart, 2010; Sheremet et al., 2016). Exhumation processes in the Greater Caucasus started in the Oligocene and reached their climax in the Miocene-Pliocene, as documented by apatite fission-track data (Vincent et al., 2007, 2011; Avdeev and Niemi, 2011). Plate reorganization has then taken place within the Arabia-Eurasia collision zone and the main plate movements have remained relatively constant over the last 5 Ma (Westaway, 1994; McQuarrie et al. 2003; Allen et al., 2004). The Rioni Basin, located between the western Greater Caucasus 227

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recent-most phase of deformation; we then illustrate the geologicalstructural data aimed at reconstructing fold geometry, and then the meso- and microtectonic data, which are useful for describing fault architecture and kinematics. Finally, we present the results of fault-slip inversion, providing the orientation of paleostress tensors, as well as their relative ages.

and the Lesser Caucasus (Fig. 1), developed mainly during the Oligocene-Miocene as a consequence of loading exerted by the AcharaTrialeti and Greater Caucasus fold-and-thrust belts (Banks et al., 1997). The study area is particularly important, also because the western Greater Caucasus developed with a predominant vergence to the south, whereas its central and eastern sectors display a double-vergence (Forte et al., 2014; Cowgill et al., 2016). The Greater Caucasus frontal folds began forming in late Eocene times; later on the deformation propagated southwards, affecting also the Rioni and Kura forelands (Banks et al., 1997). The anticlines of the northern section of the Rioni Basin apparently formed from mid-Miocene times onwards; most of their growth took place during the Meotian, followed by further, although minor growth during the Pontian-Recent (Banks et al., 1997). According to GPS data, the present convergence rate between the western Greater Caucasus and the Lesser Caucasus, across the Rioni Basin, is about 4 mm/yr (Reilinger et al., 2006). The northern portion of the Rioni Basin is characterised by a thinskinned tectonic style marked by thrusts showing a curved geometry in plan view (Banks et al., 1997; Adamia et al., 2010, 2011b). Ramp anticlines developed over south-vergent thrusts that detach and flatten along the Upper Jurassic evaporites of the Rioni foreland basin, clearly related to compression in the Greater Caucasus (Banks et al., 1997; Adamia et al., 2010).

3.1. Geomorphology The Rioni Basin area is marked by a flat geometry all along the coast of Black Sea and towards the Lesser Caucasus, with altitudes mostly < 10 m a.s.l. However, the study area is characterised by a complex morphology, with elongated highs having different orientations, and rising as high as 400 m a.s.l. (Fig. 3). The higher hills in the area coincide with structural prominent highs that coincide with the fold structures surrounding the Rioni Basin uplifted area. The interior of the uplifted area is characterised by deeply dissected alluvial fans, by a number of river terraces lying at different altitudes, as well as by several, major erosional features. The northern periphery of the study area is characterised by the Greater Caucasus mountain front, here represented by a series of planar surfaces mostly sloping to the southwest, and locally to the southeast in the western part of the mountain front. These features correspond to outcrops of steeply-dipping, substrate carbonatic layers. The elongated highs along the boundaries of the uplifted area, trend NNW-SSE in its northwestern part, NW-SE along its southwestern part, and SW-NE along its eastern margin (Fig. 3). Such topographic highs control the flow of ancient rivers; the latter, in fact, are deviated in correspondence of most of the highs (red segments in Fig. 3). One outstanding evidence of interaction between the rivers and the elongated highs can be clearly observed at a location, where a river managed to cut through an elongated high (e.g. location “A” in Fig. 3); in this case, the river got deeply entrenched, forming tens-of-meters deep gorges. This over-erosion can be explained only as evidence of hydrographic instability following intense, gradual local uplift. The latter is also attested by fluvial conglomerates of Pliocene-Quaternary age that can be found as high as 350–400 m a.s.l., thus 100–200 m above the surrounding valley floor. River terraces are particularly developed along the Enguri Valley, in the westernmost part of the elevated area. Most, if not all, are asymmetric terraces, in that they developed only along the eastern side of the river valley. As already documented by Tibaldi et al. (2017a), they were originated by eastward-directed, increasing uplift. It is worth noting that several small landslides are aligned along the southern front of the hills bordering the uplifted area, which correspond to the frontal fold limbs (e.g. Fig. 4A). Similarly, a huge landslide developed along the Greater Caucasus mountain front. This landslide contains large toreva blocks in the interior of the deposit (Fig. 4B), suggesting that it represents a deep-seated slope gravity deformation. Based on the observed stratigraphic relations, the debris avalanche deposit should be Quaternary in age.

2.2. Stratigraphy of the Rioni Basin The oldest lithostratigraphic formation exposed in the study area is composed of Lower Jurassic (about 1500 m thick) sandstones and shales, as shown in the geological map of Fig. 2. Middle Jurassic deposits are represented by Bajocian tuffs and turbidities with rare bands of calk-alkaline andesite-basalts, and by Bathonian lacustrine, coalbearing, sandy-argillaceous rocks (thickness about 2500 m). Upper Jurassic deposits are composed of clastic deposits, evaporites and basalts (about 500 m thick). A transgressive surface separates the Upper Jurassic sequence from Lower Cretaceous conglomerates at the base, followed by turbidities, dolomites, limestones, organogenic limestones and marls (thickness 350–400 m) (Adamia et al., 2011a). Above there are Upper Cretaceous, Paleocene and Eocene deposits mainly made of neritic organogenic limestones, marls and volcanogenic rocks (thickness locally > 2000 m) (Adamia et al., 2011a). The Oligocene-lower Miocene (Maykopian) series is mostly made of alternating gypsiferous clays and sandstones (thickness 800–900 m) (Banks et al., 1997; Jones and Simmons, 1997; Adamia et al., 2010). During the development of folds in the Rioni Basin, there was the deposition of syntectonic strata dating back to the Oligocene, the middle-late Miocene (Sarmatian and Meotian-Pontian), the Pliocene (Cimmerian and Kuyalnikian) and the Pleistocene (Gurian), which are characterised by onlap on fold limbs and thickening outwards, as observed at the Tsaishi fold (Tibaldi et al., 2017b). Strata geometries suggest preliminary uplifting of the frontal part of the fold in the Oligocene, followed by major folding phases along the whole structure since mid-Miocene times. The rock succession is represented by shallow marine and continental, predominantly terrigenous clastic deposits, conglomerates, sandstones, mudstones, claystones, sandy clays, clays and rare shell-beds. The total thickness of the syntectonic succession is 1500–2000 m (Banks et al., 1997; Adamia et al., 2010). Finally, part of the studied area is covered by late Quaternary fluvial deposits that give rise to a number of terraces in the interior of the uplifted area, whereas to the south lies a flat plain, made of anastomized river beds.

3.2. Structural geology of folds Here, we first briefly provide a few details on the Tsaishi fold, already documented in Tibaldi et al. (2017b), and then we proceed to describe the other folds in the arcuate belt. The Tsaishi fold, located in the southwestern part of the Rioni Basin uplift (Fig. 5), is a south-vergent anticline with sub-vertical strata to the southwest at its frontal limb, locally overturned, and shallower-dipping strata (dip 20-–40°) to the northeast, along its backlimb. From west to east, the hinge line trends N156°, N144°, N100°, N90°, and finally N126° near the eastern, conical termination, showing a sinuous shape in plan view along a total length of 29.4 km. The northwestern, apparent termination of the fold does not actually have a conical geometry, suggesting that this is an erosional feature and that the anticline originally extended further to

3. New field data In the next sections, we provide the results of the field structural research we conducted in the study area; we first describe our geomorphological observations, aimed at highlighting evidence of the 228

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Fig. 2. Map of the study area showing the main lithostratigraphic units (modified after Djanelidze and Kandelaki, 1956; Adamia and Gujabidze, 2004, and personal observations).

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Fig. 3. Digital Elevation Model of Rioni Basin uplifted area, with indication of the main geomorphological features. In blue the main rivers; in red the river segments that are offset in correspondence of the margin of the Rioni Basin uplift. “A” is a locality discussed in the text. DEM is from ALOS Global Digital Surface Model “ALOS World 3D - 30 m” (AW3D30, Tadono et al., 2014, 2016; Takaku et al., 2014, 2016). (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)

whereas in the opposite direction the fold limbs are truncated by river erosion. Philip et al. (1989) suggested that this fold is offset by a NNESSW-striking left-lateral, strike-slip fault that prolongs northward into the Enguri Valley, though there is no direct evidence of this fault plane. Northeast of the Gali anticline, another parallel syncline and an anticline can be observed. East of the Tsaishi fold, the NE-SW-trending topographic high is actually composed of a number of folds with ENE-WSW trending hinge lines (Philip et al., 1989) (Fig. 5). They comprise two major anticlines (Senaki and Martvili) and another four, minor folds, arranged according to an en-échelon, left-stepping geometry, within a N45°-trending corridor. The Senaki anticline is characterised by a southern vergence with

the northwest. Deposits as old as the Cretaceous-late Miocene are intensely deformed; the folding involves, although to a lesser degree, also younger deposits of Pliocene age; locally, slightly tilted deposits of Quaternary age have also been found (Tibaldi et al., 2017b). North of the Tsaishi fold, there is a flat-lying area that corresponds to an alluvial valley generated by deposition from the Enguri River. In the northern portion of this plain, there are some scattered outcrops made of Upper Cretaceous-Lower Paleocene (Turonian-Danian) carbonatic rocks, where it was possible to observe the presence of a NNWSSE-trending anticline (Fig. 5). This becomes clearer further north, where a main anticline with the same orientation, here named Gali fold, crops out with continuity. This fold gradually fades out northward,

Fig. 4. (A) Example of one of the several landslides aligned along the southern front of the hills bordering the uplifted area, which correspond to the frontal fold limbs. (B) Detail of a toreva block found within the deposit of the huge Quaternary landslide developed along the Greater Caucasus mountain front, here interpreted as a deep-seated slope gravity deformation. Locations in Fig. 3.

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Fig. 5. Map of main structures of Rioni Basin uplifted area, and fold strata dips plotted as lower hemisphere, Schmidt’s stereograms. The thick, continuous white lines show the kinematic trajectories, obtained as average orientations of contraction axes, calculated based on the orientation of hinge lines of folds. The latter, in turn, were reconstructed by processing field measurements of fold limb attitudes (numbers of measures given in the stereograms). White dashed lines show average fold-trend axes. Dashed black lines in the stereograms provide the horizontal projection of the main fold axis orientation. Black dip symbols provide the average attitude of strata grouped as inclination ranges.

trending N-S to NNW-SSE in the western part, NW-SE in the central part, and WSW-ENE to SW-NE in the eastern part. Thus, the folds in the interiors of the uplifted area are parallel to sub-parallel to the folds along its borders.

fold limb attitudes that indicate a N78°-trending hinge line in its western part and a N63°-trending hinge line in its eastern part. The total length of the fold is 12.2 km. Deposits dating back to the Cretaceousmiddle Miocene are involved, whereas younger deposits are not observed. To the NE is the Martvili anticline, with a hinge line trending N67° near its SW termination, N33° in its central part (whereby the scattered outcrops do not enable gaining insight into the fold geometry), and N72° at its NE termination. If we assume that the western and eastern segments of the fold are linked, its total length is 21.2 km. The different strata attitude at the fold limbs indicates a vergence to the south. Units dating back to the Paleocene and late Miocene are involved; younger deposits do not crop out. Between these two main anticlines, there are another two minor synclines and one anticline. The Gakhomela syncline, which displays the best outcrops, has a 7.7-kmlong hinge line, trending N71°. North of these folds, in the interior of the Rioni Basin uplifted area, younger deposits of Miocene-Pliocene-Quaternary age crop out; although they are mostly from horizontal to sub-horizontal, at several localities they show diverging or converging strata dips that indicate the presence of further folds (Fig. 5). These are broad synclines and anticlines, with limbs made of layers dipping 10°-–30°, and hinge lines

3.3. Fault geometry A total of 248 faults, bearing slickenside lineations, were studied in the field; their locations were recorded by way of GPS devices, and their geometric and kinematic characteristics were assessed. The faults were studied in correspondence of the main anticlines and synclines, because here there are more outcrops than in the flatter areas located further north, i.e. in the interior of the uplifted area. The faulted deposits are mainly carbonatic rocks of Cretaceous and Paleocene age and, in very few cases, of Eocene age. Other faults were measured in the northern sector of the study area, in correspondence of outcrops of Cretaceous carbonatic rocks along the foothills of the Greater Caucasus. Fault scarps locally affect younger deposits too, as recent as the Quaternary, but poor outcrop conditions prevent from locating suitable fault planes with slickensides. Faults with very different strikes, as well as belonging to all 231

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faults are characterised by low dip angles (most of which shallower than 60°) and are grouped in clusters, with fault azimuth values ranging N20°–N60°, N95°–N105° and N130°–N150°. The second cluster is characterised by the shallowest dip angles, as shallow as 20°. Normal faults are characterised by fault azimuths ranging from N60° to N130° and dip angles always steeper than 50°. Right and left-lateral strike-slip faults are marked by heterogeneous azimuth values: those with fault azimuths in the N0°–N100° range are characterised by fault dips steeper than 55° (with the exception of two left-lateral faults, whose fault dips are between 30 and 40°), whereas strike-slip faults with azimuths in the N101° to N180° range have dip angles as shallow as 35°. The observed dip angles, which are shallow for transcurrent faults, might suggest post-faulting tilting, as will be discussed later on. 3.4. Microtectonic data We collected hundreds of microtectonic data at several, homogeneously distributed measurement sites, along the arcuate belt of folds, and at fewer sites along the northern portion of the Rioni Basin at the southern Greater Caucasus front. On the other hand, in the interior of the uplifted area we could not gather any microtectonic data, owing to to the very poor outcropping conditions, affected by a dense and diffuse vegetation cover. Microtectonic data comprise slip indicators along fault planes (tectogliphes) (Fig. 7A–E), as well as axes of tectonic stylolites (Fig. 7F). The most common tectogliphes which we assessed are striae, Riedel and Riedel1 microfractures, crystal fibres, steps and slickenside flutes (following nomenclature by Petit, 1987). Fault slip data indicate that the whole range of possible kinematics is represented in the study area: right-lateral strike-slip faults (103) are the most frequent (Fig. 8A), followed by left-lateral strike-slip faults (72) (Fig. 8B), reverse faults (62) (Fig. 8C) and normal faults (11) (Fig. 8D). Right-lateral strike-slip faults tend to cluster in three main sets, striking NNE-SSW, NNW-SSE and WSW-ENE. Left-lateral strike-slip faults tend to cluster in three main sets, striking NW-SE, WSW-ENE and NNE-SSW, in order of decreasing frequency. The occurrence of these sets indicates that more than one phase of transcurrent deformation took place, as indicated also by the systematic crosscutting relationships among faults with different orientations observed in the field. Reverse faults belong to three main sets: The most frequent is made of fault planes dipping to the SE, followed by those dipping to the SW and NW. Faults dipping to the NW and to the SE crosscut each other without a systematic time-relation and, therefore, they can be regarded as two conjugated sets belonging to the same phase of deformation. Normal faults are very limited in number and strike NW-SE; they are made of two sets having opposite dips that do not offset each other with any recognisable time-relation; thus, they form a conjugate system. As will be reported in the following section, the orientation of faults of the same age and kinematics can slightly change in space, as shown by the distribution of the paleostresses. Fig. 6. (A) Fault strike frequency. (B) Fault dip frequency. (C) Graph of fault strike versus fault dip. Blue and green circles represent right and left-lateral strike-slip faults, respectively; black squares represent normal faults; red diamonds represent reverse faults. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)

3.5. Paleostress tensors 3.5.1. Method Paleostress tensors were computed through the inversion of faultslip data related to populations of structures of the same age. The relative age of the faults was assessed based on: (i) crosscutting relationships among striae, in the case of occurrence of more than one generation of slickensides on any fault plane (e.g. Fig. 7A); (ii) the statistical persistence of a given set of planes with respect to others having a different orientation in the same rock mass (e.g. Fig. 7B); (iii) the offset relationships among fault planes (e.g. Fig. 7C–D); (iv) the kinematic compatibility; and (v) the age of the offset lithostratigraphic units. In order to improve the dating of the various fault phases, we placed particular emphasis on the detailed age of the geological units affected by faults, also keeping into account information published in the literature, geological maps (Djanelidze and Kandelaki, 1956;

kinematic types, are present, giving rise to a structural pattern that is more complicated than previously documented for the same area by Philip et al. (1989). The graph of Fig. 6A shows the frequency of fieldcollected fault strikes in the whole study area, which is given by faults striking mostly N20-–40°, N90-–110° and N120-–130°, though minor sets are present also in the remaining azimuth ranges. The most frequent fault dip is represented by vertical to sub-vertical (> 70°) planes, followed by the 51-–60° range, and then by 41–50°, 61–70°, 31–40° and 21–30° in decreasing order of frequency (Fig. 6B). The graph of Fig. 6C shows a plot of fault strikes versus fault dip for the whole studied population, broken down by kinematics. Reverse 232

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Fig. 7. Examples of kinematic indicators and fault time-relations found in the study area. (A) Two generations of striae cross-cut each other, allowing to assess the younger slickenside lineation (site 254a in Fig. 12). (B) The younger faults found at any outcrop are represented by major, continuous right-lateral strike-slip planes, which offset all other faults (site 58b). (C) and (D) Photo and sketch of a fault plane (dark grey) offsetting another fault (light grey) according to a right-lateral strike-slip component of motion (site 254b). (E) Left-lateral strike-slip fault with step-like crystal fibers (syn-cinematic crystallization fibers) (site 214). (F) Horizontal tectonic stylolite peaks (site 73).

Adamia and Gujabidze, 2004) and personal field observations. The faulted lithostratigraphic units range from Middle Jurassic to Cretaceous, Paleocene and Eocene times. The faults of the same age were processed by way of the SG2PS software (Structural Geology to Post Script Converter - http://www. sg2ps.eu; Sasvári and Baharev, 2014), which was used for reconstructing the stress tensor. The software enables performing paleostress inversion using methods by Turner (1953), Sprang (1972), Michael (1984), Angelier (1990), Fry (1999), Shan et al. (2004), and Mostafa (2005). We carried out paleostress analyses by using the INVD direct inversion method (Angelier, 1990), capable of using fault slip data for determining the stress tensor with four degrees of freedom (orientation of the principal stress axes and the ratio of principal stress differences). The software calculates also a misfit vector “ν” between the measured and calculated shear vector, and minimizes its length, for at least four different striated faults of the same age. The length of vector ‘v’ is expressed as a function of the plane normal vector, the slickenside lineation, the assumed (a priori) shear stress vector length λ and the members of stress tensor. Furthermore, this method enables calculating as great as possible shear stress magnitudes, in such a way as to overcome cohesion and friction on fault planes. The software output is composed of stereograms containing fault planes associated with their kinematics and the orientation of the principal stress axes, namely the greatest principal stress (σ1), the intermediate principal stress (σ2) and the least principal stress (σ3). Associated with each

stereogram is a colour-bar legend, which, by way of a solid white line, indicates the tectonic regime, ranging from red, which represents compression, to green, representing transcurrence, to blue, representing extension. Faults were divided also based on their relations to folding: Those which developed prior to folding were back-tilted and re-processed with the purpose of calculating the original geometry of stress axes (e.g. Saintot and Angelier, 2002). As values for the back-rotation of fault planes, we used the dip and dip angle of the fold limb (or tilted strata at the Greater Caucasus front), which we measured in the field at the same sites where we collected fault-slip data. Table 1 shows the results of all stress tensors as resulting from present-day fault geometry, whereas Table 2 shows only the stress tensors of those sites where back-tilting was applied. 3.5.2. Results for fault-slip inversion The computed stress tensors based on present-day fault geometry are displayed in Figs. 9–11 (and Table 1). The data are represented as Schmidt’s stereograms, lower hemisphere, with ciclographic projections of planes and kinematic data. The computed stresses are represented with diverging arrows that provide the orientation of σ3, and converging arrows that provide the orientation of σ1. The plots are listed in agreement with the number of site measurement, whereas their location and the orientation of the stress axes computed after back-tilting are shown in Fig. 12. 233

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solutions (Tsereteli et al., 2016) as well, and thus the stress field trajectories were obtained as average orientation of the conterminous σ1 axes and earthquake P (compression) and T (tension) axes. In regard to extensional deformation, we also measured veins, extensional joints and crystal fibers. However, we did not subdivide the extensional structures in terms of age, because they are very limited in number and of uncertain time attribution; nevertheless, they are considered in the discussion. During Phase I (Fig. 12A), the σ1 trajectories trend NNW-SSE in the northern part of the study area, and they tend to gradually rotate clockwise in the western part of the study area. As a whole, σ1 trajectories tend to be normal to fold hinge lines. This phase is mostly represented by the activity of transcurrent faults and subordinate reverse faults. All stress axes at these measurement sites were inclined (before restoration), and thus consistent with a pre-folding age of this phase of deformation. The age of the involved lithostratigraphic units ranges from the Middle Jurassic, to Early and Late Cretaceous, until the beginning of Palaeocene (Danian). As regards Phase II (Fig. 12B), σ1 trajectories trend, on average, WNW-ESE in the western sector of the study area, and preferentially EW in the eastern sector. This phase is represented by the activity of transcurrent faults with horizontal σ1 and σ3. Compression is from parallel to oblique to the hinge lines of the folds. Also these stress axes were inclined (before restoration), and this suggests that this phase of deformation developed prior to folding. The age of the involved lithostratigraphic units ranges from the Cretaceous to Palaeocene and in one case (site 31) also the Eocene. During Phase III (Fig. 12C), σ1 trajectories trend NE-SW in the western part of the study area, NNE-SSW in the central part and NNWSSE in the eastern part, resulting in a southward-opening fan geometry. Compression axes are oblique to the fold hinge lines of the eastern part of the arcuate fold belt, where a general left-lateral shear is expected, and this is consistent with the left-stepping, en-échelon arrangement of all these folds. This phase is represented essentially by transcurrent faults with horizontal σ1 and σ3. All computed stress tensors are characterised by vertical and horizontal stress axes, consistent with this phase being successive to the main folding. The age of the involved lithostratigraphic units ranges from the Cretaceous to Palaeocene. Phase IV (Fig. 12D) shows dominant N-S to NNE-SSW σ1 trajectories, obtained by integrating inversion of fault-slip data with earthquake focal mechanism solutions. The latter are represented by reverse motions along about E-W striking fault planes, located in the northern sector of the study area; the southern sector, in contrast, is characterised by transcurrent faults. All computed stress axes are originally un-rotated and thus correspond to post-folding deformation. The age of the involved lithostratigraphic units ranges from the Cretaceous to Palaeocene. Some fault scarps that involve the rock succession up to the Quaternary are oriented consistently with this state of stress, although they do not show slickensides.

Fig. 8. Graphs of field-surveyed faults subdivided in terms of their kinematics: Left column shows stereographic plots, the right column rose diagrams of fault strikes. (A) Right-lateral strike-slip kinematics; (B) left-lateral strike-slip kinematics, (C) reverse kinematics, (D) normal kinematics.

Compression and extension directions are not homogeneous and several ranges are present; the most frequent stress tensor is represented by horizontal σ3 and σ1 that correspond to transcurrent deformation phases. Less frequent are stress tensors with horizontal σ2 and σ1 (reverse faulting) and still less frequent are stress tensors with horizontal σ2 and σ3 (normal faulting). The σ1 axes related to transcurrent deformation tend to cluster into NE-SW, a NNE-SSW, and a WNW-ESE to E-W trends. The σ1 axes associated with reverse deformation tend to range from NNW-SSE to NNE-SSW. Normal faults are associated with a NE-SW-trending σ3. The orientations of σ1 and σ3 axes are represented in Fig. 12 after back-tilting (results in Table 2) and they are divided into four main phases of deformation, based on the relative age of the brittle structures used for slip data inversion, as previously discussed. Along with field age relations between the various features, we also considered prefolding or post-folding brittle deformation. For each of the four phases of deformation, we reconstructed the stress field as σ1 trajectories. These are calculated as average orientation of the conterminous σ1 axes. For the younger phase, we used the published focal mechanism

3.5.3. Results for stylolites We measured about 40 stylolite peaks at outcrops in different locations along the arcuate fold belt (Fig. 11). Only well-developed peaks and with axes ranging from horizontal to gentle-dipping were considered and subjected to measurement. These types of stylolites are tectonic in origin; we also identified the stylolites produced by lithostatic loading, marked by normal-to-bedding peaks. These kinds of stylolites, in the study area, occur only in the carbonatic rocks that lie in the lower section of the outcropping rock succession, Cretaceous in age. We regard the perfectly horizontal peaks as belonging to the younger deformation phase, which postdates folding. On the other hand, we regard the inclined stylolite peak sets as having been rotated and coeval to folding. We back-rotated inclined peaks in order to restore them to their original position. As the various sets show different orientations, we assigned them to the relevant deformation phases based on their folding or post-folding age, and on the similarity to the σ1 calculated by 234

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Table 1 Columns show: site number, site location (Lat and Lon), number of data used during the inversion process, plunge and dip of resulting σ1, σ2 and σ3, ratio (R (Φ)) between the differences of the principal stress eigenvalues, (σ2 − σ3)/(σ1 − σ3), Av Misfit Angle = average angle between computed shear stress and slip vector, resulting tectonic regime. Site

Lat (dd°)

Lon (dd°)

N. of data

σ1 (plg/dip)

σ2 (plg/dip)

σ3 (plg/dip)

R (ϕ)

Av misfit angle

Tectonic regime

1 2 67 73 145 167 245 103a 103b 159–160a 159–160b 167–170 249–253 254a 254b 254c 254d 254e 66a 31 36 37 50 58a 58b 68 184a 184b 225 220 214 192 200

42.76 42.39 42.41 42.39 42.37 42.29 42.42 42.36 42.36 42.43 42.43 42.42 42.41 42.42 42.42 42.42 42.42 42.42 42.41 42.32 42.30 42.30 42.28 42.46 42.46 42.36 42.66 42.66 42.46 42.46 42.46 42.64 42.64

42.03 41.83 41.80 41.82 42.19 42.10 41.80 42.19 42.19 41.81 41.81 41.83 41.80 41.81 41.81 41.81 41.81 41.81 41.80 42.14 42.13 42.13 42.07 42.37 42.37 41.87 42.22 42.22 42.37 42.37 42.38 42.20 42.20

6.00 6.00 7.00 6.00 10.00 6.00 7.00 5.00 7.00 6.00 5.00 5.00 5.00 6.00 4.00 7.00 5.00 13.00 5.00 5.00 9.00 11.00 6.00 20.00 5.00 7.00 5.00 4.00 5.00 4.00 4.00 4.00 7.00

352/07 013/00 193/13 43/17 193/11 197/13 186/39 346/17 314/05 300/15 232/02 017/34 042/02 342/17 190/17 015/03 235/22 125/01 118/39 259/01 291/55 170/21 192/11 265/22 339/15 110/11 190/10 355/32 281/14 102/17 245/06 167/19 148/31

085/29 103/00 338/74 141/25 058/75 288/04 033/47 083/21 224/01 204/22 322/05 283/07 311/07 083/32 071/58 155/86 132/29 216/50 209/01 167/60 125/34 303/60 350/78 52/65 136/74 292/79 284/22 091/10 057/71 347/54 342/48 041/60 251/21

250/60 207/90 101/09 283/59 285/10 033/76 287/14 221/63 125/85 061/63 119/85 183/56 144/83 228/53 288/26 285/03 356/52 033/40 300/51 350/30 30/07 72/22 101/04 170/12 247/06 200/00 076/65 196/57 187/13 202/31 149/41 266/23 010/51

0.52 0.48 0.32 0.65 0.34 0.87 0.684 0.17 0.74 0.131 0.278 0.903 0.781 0.411 0.556 0.404 0.509 0.154 0.09 0.02 0.644 0.4 0.095 0.257 0.26 0.425 0.45 0.54 0.44 0.20 0.03 0.85 0.48

1.40 2.50 13.30 2.10 3.50 4.60 15.6 2.10 1.70 2.4 3.6 15.6 10.8 5.9 9.1 13.8 3.7 12.5 5.70 3.70 27.50 9.30 2.30 4.70 0.90 2.10 7.40 9.00 3.80 4.10 2.80 28.90 11.00

PURE COMPRESSIVE PURE COMPRESSIVE PURE STRIKE-SLIP PURE COMPRESSIVE PURE STRIKE-SLIP RADIAL COMPRESSIVE PURE STRIKE-SLIP TRANSPRESSIVE PURE COMPRESSIVE TRANSPRESSIVE PURE COMPRESSIVE RADIAL COMPRESSIVE RADIAL COMPRESSIVE PURE COMPRESSIVE PURE STRIKE-SLIP PURE STRIKE-SLIP PURE COMPRESSIVE TRANSPRESSIVE TRANSPRESSIVE TRANSPRESSIVE PURE EXTENSIVE PURE STRIKE-SLIP TRANSPRESSIVE PURE STRIKE-SLIP PURE STRIKE-SLIP PURE STRIKE-SLIP PURE COMPRESSIVE PURE COMPRESSIVE PURE STRIKE-SLIP TRANSPRESSIVE TRANSPRESSIVE TRASTENSIVE PURE COMPRESSIVE

direction normal to the plane of extensional fissures, from crystal fibres axes growth in veins, and from inversion of normal fault-slip data. Extension produced stretching along the arc, with orientation parallel to fold hinge lines in the western part of the belt, and from parallel to subparallel to the folds in the eastern part. The amount of stretching appears to be quite limited, and it is not worthy of quantification, due to scattered observations at poor-quality outcrops. For the same reason, it was impossible to establish a time-relation between extensional structures and the various phases of compressional deformation. Only in a few cases it was possible to conclude that extensional fractures are younger than the other structures, though at site 36 the normal faults resulted to predate right-lateral strike-slip faults.

way of fault slip inversion. The oldest peaks trend NNE-SSW (Fig. 12A), hence indicating a σ1 consistent with the stress field derived from fault slip inversion for this phase in the southwestern part of the study area. These peaks were measured along stylolitic planes that are offset by some fault sets, a fact that confirms they belong to the oldest stress phase. The younger peaks trend from WSW-ENE to WNW-ESE, consistently with the σ1 trajectories of Phase II (Fig. 12B). The stylolites measured at site 68b are of difficult age attribution, thus they were tentatively assigned to Phase I.

3.5.4. Extensional structures Measurement sites and the resulting extensional directions (diverging arrows) are shown in Fig. 13, whereas single structures are shown in Figs. 9 and 11. Extensional directions were obtained by assessing the Table 2 Paleostress tensors calculated after back-tilting of faults. Columns as in Table 1. Site

Lat (dd°)

Lon (dd°)

N. of data

σ1 (plg/dip)

σ2 (plg/dip)

σ3 (plg/dip)

R (ϕ)

Av misfit angle

Tectonic regime

1 73 167 245 103a 160 167 254a 254b 254d 254e 66a 31 58a

42.76 42.39 42.29 42.42 42.36 42.43 42.42 42.42 42.42 42.42 42.42 42.41 42.32 42.46

42.03 41.82 42.10 41.80 42.19 41.81 41.83 41.81 41.81 41.81 41.81 41.80 42.14 42.37

6.00 6.00 6.00 7.00 5.00 6.00 5.00 6.00 4.00 5.00 13.00 5.00 5.00 20.00

68/59 193/25 227/09 200/09 16/55 162/09 277/09 333/12 012/17 055/07 128/13 161/23 094/19 249/03

281/26 58/57 183/25 319/71 272/10 017/79 183/25 075/44 201/73 150/36 224/23 053/37 342/46 157/31

184/14 293/21 026/63 107/16 176/33 253/06 26/63 232/43 103/02 315/53 011/63 276/45 199/38 344/59

0.868 0.570 0.829 0.564 0.647 0.755 0.829 0.352 0.440 0.501 0.194 0.447 0.167 0.129

1.2 3.3 4.7 17.2 2.2 2.5 4.7 6.2 8.1 3.2 14 5.1 3.6 4.5

TRANSTENSIVE PURE STRIKE-SLIP RADIAL COMPRESSIVE PURE STRIKE-SLIP TRANSPRESSIVE TRANSTENSIVE RADIAL COMPRESSIVE PURE STRIKE-SLIP PURE STRIKE-SLIP PURE COMPRESSIVE TRANSPRESSIVE PURE COMPRESSIVE TRANSPRESSIVE TRANSPRESSIVE

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Fig. 9. Schmidt’s stereograms, lower hemisphere, of faults grouped by coeval sets, showing fault-slip vectors (small arrows along great circles), calculated horizontal orientation of the greatest principal stress σ1 (converging black arrows) and least principal stress σ3 (diverging white arrows) by inversion of slip vectors, and location of principal stress axes (σ1 red diamond, σ3 blue diamond, σ2 green diamond). White rectangle in coloured bar indicates the tectonic regime. Numbers refer to localities shown in Fig. 12. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)

4. Discussion

stress axes related to folding and we performed the needed back-restoration. However, we cannot rule out the occurrence of some block rotation along vertical axes, possibly related to the development of a shear type of deformation. We need to stress that checking this possibility is not feasible for the time being, because no paleomagnetic data are available in the area as far as we know.

Our geological-structural analyses allowed us to reconstruct the various paleostress fields that have affected the Rioni Basin area over time. By analysing the brittle structures surveyed in the field, coupled with the calculation of the paleostress tensors, we have also been able to infer relations with fold geometry. We considered the rotation of the 236

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Fig. 10. Same legend as in Fig. 9.

4.1.1. Phase 1 The first deformation phase is characterised by σ1 trajectories that are perpendicular to the main Greater Caucasus mountain belt and to most folds in the Rioni Basin. Moreover, the brittle structures produced by this stress field were rotated by fold growth. In order to come up with a possible absolute dating for this first phase, we need to take into consideration that the stratigraphic relations among the syn-depositional strata onlapping the Tsaishi fold limb suggest that folding started in Oligocene times and went on with a peak in middle Miocene times (Banks et al., 1997; Tibaldi et al., 2017b). As our Phase 1 affects rocks of Middle Jurassic, Cretaceous and Paleocene age, and it predates Oligocene-Miocene folding, we can conclude that it took place in the

4.1. Deformation phases Four main deformation phases were identified in the studied rock succession. These are fewer than the number of events recorded farther west, in the Russian Greater Caucasus by Saintot and Angelier (2002). This can be explained by keeping into account that these authors investigated a much larger area and a more complete rock succession. Moreover, our study area is not related to the inner core of the Greater Caucasus, but to the uplifted area located south of the main mountain belt. Thus, it is possible that our study area was affected by fewer deformation events than the main mountain belt.

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Fig. 11. For the first six stereograms, the legend is the same as in Fig. 9. Grey divergent arrows with great circles indicate σ3 obtained by fissures and extensional joints plotted in Schmidt’s stereograms, lower hemisphere, as planes. Grey divergent arrows for blue rose diagrams indicate σ3 obtained by veins. Converging black arrows for rose diagrams indicate σ1 axis obtained by populations of stylolite peaks. Site 113b give σ3 obtained by axes of a population of crystal fibers grew in veins. Numbers refer to localities shown in Fig. 12. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)

may indicate rotation along a vertical axis successive to the formation of the faults from Phase 1. In this case, rotation was clockwise along the folds of the western sector of the arcuate belt.

Eocene. A pre-folding phase of faulting was recognized in the western Greater Caucasus (NW of our study area) by Saintot and Angelier (2002); this was characterised by a first strike-slip faulting event, followed by a reverse faulting event. Both were caused by a NE-SW to NNE-SSW σ1 and were dated to the late Eocene by the same authors. Although an even older phase of deformation (Paleocene-early Eocene) was observed in the Crimea mountains by Sheremet et al. (2017), we are inclined to suggest that our Phase 1 may coincide with the late Eocene phase recognized by Saintot and Angelier (2002), based on the similar orientation of the σ1, fault kinematics, estimated time interval and proximity of the studied areas. In the western Greater Caucasus, this phase played a prominent role in the development of folding and thrusting (Milanovsky and Khain, 1963; Grigor’yants et al., 1967; Milanovsky et al., 1984; Muratov et al., 1984; Mikhailov et al., 1999), with the formation of southwest-vergent thrusts under a NNE-SSW to NE-SW directed σ1 (Khain, 1984; Giorgobiani and Zakaraya, 1989), involving also the offshore region of the eastern Black Sea (Finetti et al., 1988; Terekhov and Shimkus, 1989; Robinson et al., 1996). However, it is worth highlighting that, in our study area, folding did not develop at this stage and regional compression was absorbed in the form of diffuse brittle deformation. The fan-shaped arrangement of σ1 trajectories in the western part

4.1.2. Phase 2 Phase 2 is characterised by strike-slip faults produced by an WNWESE to E-W trending σ1, and pre-dates folding. These kinematics and stress trajectories coincide with the features of the tectonic phase recognized in the western Russian Greater Caucasus by Saintot and Angelier (2002). The authors dated this phase to the beginning of Miocene, and suggested that, in their study area, it postdates folding. In regard to the latter observation, we wish to highlight that in the Rioni Basin, in general, the main fold growth took place in mid-Miocene times (Banks et al., 1997), i.e. later than in the western Russian Greater Caucasus. However, detailed seismic stratigraphic analyses in correspondence of the Tsaishi fold carried out by Tibaldi et al. (2017b), suggest that minor uplift/folding events occurred in the Oligocene and beginning of the middle Miocene, followed by further folding events. As a consequence of the above, and taking into consideration the Cretaceous-Paleocene and Eocene age of the faulted rocks, an early Miocene age for Phase 2 fits with the rotation of the computed stress axes due to successive folding, as observed in our area for the relative stress tensor solutions. 238

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Fig. 12. Calculated stress orientation and trajectories of σ1 referring to: (A) the oldest tectonic phase, represented by reverse faulting (converging arrows give orientation of σ1) and transcurrent faulting (converging arrows give orientation of σ1 and diverging arrows give orientation of σ3). (B) and (C) The second and third tectonic phase, mainly represented by transcurrent faulting. (D) The youngest tectonic phase, showing also P-axes of four main crustal earthquakes (from Tsereteli et al., 2016). Numbers refer to full data sets shown in Figs. 9–11.

developed, with predominantly right-lateral strike-slip displacements. The origin of the rotation of the σ1 from a NNE-SSW trend during Phase 1 to a WNW-ESE and an E-W orientation for Phase 2 is difficult to account for. In early Miocene times (> 20 Ma), the convergence of the Arabian Plate toward Eurasia occurred along a NW-SE direction, followed by the development of compressional structures in the whole Black Sea-Caucasus region (Savostin et al., 1986; Zonenshain et al., 1990). Saintot and Angelier (2002) suggested that, during these times, the Greater Caucasus central zone may have acted as a nucleus for

A similar phase of deformation was recognized also in the northwestern Greater Caucasus by Shardanov and Peklo (1959) and Kopp and Shcherba (1998), and in the western Greater Caucasus by Saintot and Angelier (2002). In that area, it was characterised by a WNW–ESE transpressional paleostress field, named “Save tectonic phase” and dated to early Miocene times. In particular, Rastsvetaev et al. (2010) found a widespread system of WNW-ESE-striking right-lateral strikeslip faults in the north-western Caucasus. In the studied area, under Phase 2 stress field, previously existing faults and new faults may have

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Fig. 13. Locations of measurement and extensional directions (diverging orange arrows) obtained from extensional fissures, normal faults, and crystal fibres axes. Note that extension is from parallel to subparallel to the local orientation of the arcuate belt. Fold axes and main reverse faults are in grey. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)

4.1.3. Phase 3 In Phase 3, the stress field was dominated by σ1 trajectories trending in the range from NNE-SSW to NE-SW (Fig. 12C). As one of the three main computed stress axes is always vertical, this phase should have started immediately after the main folding phase. The climax in fold growth, followed by faulting, most likely took place during the regional deformation phase of the middle Miocene, in response to the indentation of the Arabian plate into the Caucasian syntax (Beloussov, 1940; Shardanov and Peklo, 1959; Milanovsky and Khain, 1963; Beliaevsky et al., 1961; Shcherba, 1987; Kopp, 1989; Giorgobiani and Zakaraya, 1989; Rastsvetaev, 1989; Kopp and Shcherba, 1998; Zonenshain et al., 1990; Milanovsky, 1991; Nikishin et al., 1998). This triggered main folding with hinge lines trending WNW-ESE and thrusting along WNWESE-striking planes along the Greater Caucasus belt. However, in the study area the orientation of σ1 trajectories caused left-lateral shear along the eastern portion of the arcuate belt, with the development of the deformation corridor marked by the left-stepping, en-échelon arrangement of folds. As the asymmetry of the fold limbs indicates vergence to the south, this zone can be regarded as a leftlateral transpressional shear. Within this zone, we suggest that

uplifting and contractional processes, resulting in a strong lateral contraction to the sides, i.e. the western and to the eastern Caucasus. According to these authors, the lateral push caused left-lateral shear along major faults parallel to the mountain belt. It is worth stressing that this scenario cannot be applied to our study area, because the main brittle structures from Phase 2 are represented by right-lateral strike slip faults striking mostly from E-W to NW-SE and by subordinate, left-lateral strike-slip faults striking around NNE-SSW. We suggest that the WNWESE-directed σ1 in the Rioni Basin was the effect of the combination of the main NW-SE convergence direction of the Arabian Plate toward Eurasia and the pre-existence of faults oriented parallel to sub-parallel to the Caucasus trend. This resulted in right-lateral transcurrence along the main Caucasus fault trend, and left-lateral shear along antithetic (Riedel 1) faults perpendicular to the mountain belt. We are also inclined to believe that the curved shape (in plain view) of the σ1 trajectories belonging to this phase, might be explained in terms of an anticlockwise rotation, along a vertical axis, of the tectonic blocks located in the eastern sector of the studied arcuate belt. The rotation should have taken place during the successive two deformation phases.

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the study area and that are oblique in its eastern sector, whereas Phase 4 is characterised by compression essentially oblique to the axes of folds. Therefore, we are inclined to argue that the main factor controlling arc geometry was different from the pure regional compression trajectories. We need to consider that seismic reflection sections and field geological-structural data indicate that the crust of the uplifted Rioni Basin is characterised by the widespread presence of south-vergent thrusts with a ramp-and-flat geometry (Banks et al., 1997; Tibaldi et al., 2017a, 2017b). In particular, the data provided by Tibaldi et al. (2017b) indicate that the major Tsaishi anticline is a fault-propagation fold. Since the evolution and geometry of fault-propagation folds is intimately linked to the geometry of the underlying thrust plane, we suggest that the formation of this fold belt and the development of its arcuate shape were controlled by two main factors: The margin geometry of the overthrusting block (acting as an indenter), and the block’s direction of movement relative to the orientation of its margins. Several analogue experiments studied the different geometries of an arcuate belt, obtained by varying these parameters (Zweigel, 1998; Zweigel et al., 1998; Macedo and Marshak, 1999; Lickorish et al., 2002; Marques and Cobbold, 2002). Further factors may control the final arcuate belt geometry, such as variations in the pre-deformational sedimentary thickness, whereby particularly thick sedimentary basins may result in salients (Thomas, 1977; Marshak and Wilkerson, 1992; Lawton et al., 1994; Boyer, 1995). However, the studied uplifted area belongs to the much wider Rioni Basin and, based on the current knowledge, the uplifted area does not seem to coincide with a major local thickening of the rock succession. Another factor which is worth pointing out is represented by the variations in the strength of the detachment horizon, that, in turn, may influence the width and geometry of an arcuate belt (e.g., Davis and Engelder, 1985; Luján et al., 2003). In the present case study, the limited dimension of the area and the homogeneous distribution of rock types do not indicate the presence of important variations in decollement material properties. However, previously published sections showed that the main deeper decollement surface lies in correspondence of evaporite deposits (Banks et al., 1997; Tibaldi et al., 2017a, 2017b) and thus it is not possible to rule out that some local change in the present evaporite thickness may be present. Finally, inherited weakness zones may contribute to dictating the foldand-thrust belt shape (Calignano et al., 2017), but also this possibility does not seem to apply here. In order to establish a direct comparison between analogue models and field cases, the knowledge of the motion vectors is needed. In this regard we are aware that, in the case under study, the movement direction of the overthrusting crustal block is unknown, and only 3-D restorations of an arc may provide the relevant data related to displacement directions (Hindle and Burkhard, 1999). In their absence, the most useful criteria for the distinction of arc formation modes are the reconstruction of the strain pattern with fold contraction axes, together with the evaluation of the presence of arc-parallel stretching and rotation of portions of the arc, in comparison with the general deformation geometry. We thus consider the following constraints, as observed in the field in the Rioni Basin uplifted zone: (i) the general geometry of the deformed zone; (ii) the presence of local extensional deformation with stretching parallel to the arc; (iii) the presence of a clear left-lateral transpressional shear zone along the eastern margin of the arcuate belt; (iv) possible vertical axis rotations along the eastern belt; (v) the fan-shaped pattern of the strain trajectories; (vi) the more diffuse folding in the western sector of the Rioni Basin uplift that involves a larger area than in the eastern sector. In regard to the general geometry of the studied area, the distribution of outcropping main faults and folds provides the trend-line pattern showed in Fig. 5 (white lines). By comparing this with the fundamental five types described by Macedo and Marshak (1999), it is possible to determine that the Rioni Basin salient is of the symmetric type, with a convergent trend-line pattern characterised by trend lines converging to both end points. In this pattern, the spacing between the

anticlockwise rotation along a vertical axis did take place; this is consistent with the reconstructed left-lateral transpression, with the dominance of strike-slip faults during Phase 3, and enables providing an explanation for the rotated σ1 trajectories during Phase 2. We suggest that this local anticlockwise rotation continued also later on, and this accounts for the fan shape of σ1 trajectories from Phase 3. 4.1.4. Phase 4 During Phase IV, σ1 trajectories are oriented from N-S to NNE-SSW. Horizontal stress axes of the calculated paleostresses are un-rotated, and also the stress field lines are parallel to each other. These data suggest that the faults belonging to Phase 4 were not rotated, either by folding or by wrench tectonics. These faults are the youngest that we found in the field bearing slickenside lineations, but we cannot date them exactly. However, we need to point out that a main, NW-SE striking reverse fault has been recently documented as bearing evidence of recent (Quaternary and possible Holocene) movements by Tibaldi et al. (2017b). This corresponds to the Tsaishi fault, displayed in Fig. 12D. Another major (possibly Pleistocene-Holocene) structure, recognised by Tibaldi et al. (2017a), is represented by an E-W-striking reverse fault located in the inner sector of the study area. These two structures have orientations that are compatible with the reconstructed stress field. The continuation of this stress field until the present day is attested by the four focal mechanism solutions of major earthquakes (M 3.5–4.7) that suggest reverse to transpressional slip along E-W- to WNW-ESE-striking fault planes. The recentmost stress field of the uplifted Rioni Basin is also consistent with the regional σ1 trajectories all across Georgia, as shown by Tsereteli et al. (2016), which have a N-S to NNE-SSW orientation. Also the several aligned landslides that we found in the studied area, are an indirect evidence of recent tectonic activity. We found shallow landslides and a huge deep-seated gravity slope deformation of late Quaternary-Holocene age along the southern front of the Greater Caucasus, and several shallow landslides of recent age sitting on top of any other deposits, especially along the southern limbs of the Rioni arc frontal folds. Such alignments of landslides along tectonic structures are classically regarded as evidence of gravity deformation, possibly induced by recent tectonic displacements (e.g. Tibaldi and Ferrari, 1992; Tibaldi and Pasquaré, 2008; Tibaldi et al., 1995, 2015). This scenario, pointing to ongoing deformation, is corroborated by recent GPS data and plate tectonic models, indicating that the Greater and Lesser Caucasus are currently very tectonically active (Rebai et al., 1993; Koçyiğit et al., 2001; Reilinger et al., 2006; Tan and Taymaz, 2006; Pasquaré et al., 2011). This can be explained in terms of the longlived convergence that is still going on after the continent-continent collision between the Eurasian and Africa-Arabian plates (Avagyan et al., 2010; Adamia et al., 2017). 4.2. Factors controlling arc formation The stress fields in the uplifted Rioni Basin, reconstructed by way of fault-slip inversion, show that N-S to NE-SW compression has continued during most of Tertiary-Quaternary times, with the exception of Phase 2, at a time when the stress trajectories changed. This means that this portion of the Rioni Basin gradually closed in response to a major shortening pattern, consistently with the general geometry of the Greater Caucasus. Nevertheless, as opposed to the main Caucasus mountain belt, the geometry of the uplifted Rioni Basin is characterised by a clear arcuate shape in plan view, and thus it is scientifically challenging to try explaining this difference. As can be seen in Fig. 5, fold hinge lines have two main trends, roughly perpendicular to each other. Although important differences in the orientation of σ1 trajectories took place over time, Phase 1 and 2 cannot account for the different orientation of folds, simply because these phases predate folding. Phase 3 is marked by compression trajectories that are perpendicular to fold axes only in the western part of 241

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Fig. 14. Results of analogue models mimicking the indentation of a rigid block into deformable material (sand) with different geometries of the indenter with respect to its motion direction. All plan views. (A) The indenter moves along a path perpendicular to the front; transcurrence takes place at both indenter margins and a rectilinear fold-and-thrust system develops along the front. (B) The indenter moves obliquely with symmetric margins with respect to the motion vector; an arcuate symmetric fold-and-thrust system develops. (C) Photo of asymmetric oblique indentation and (D) interpretative sketch showing the different tectonic regimes at the front (a wider fold-and-thrust system) and at the margin (a narrow shear zone with enéchelon structures). (Modified after Zweigel, 1998).

traces of first-order structures decreases towards both of the end points in the salient. Anyway, the general symmetry of trend lines as seen in plan view is complicated by the variations of tectonic regime along strike of the arcuate belt. Analogue models, in fact, show that the tectonic regime along the arcuate belt can change from pure contraction to transpression or transcurrence, depending upon different angular relations between the temporally changing movement direction of the indenter and the laterally changing orientation of the plate margin (Zweigel et al., 1998). In the case of frontal indentation, which is with movement normal to face of indenter, no orogen-parallel extension occurs along the front of the indenter (Fig. 14A), whereas high orogenparallel extension develops around the corner of the indenter. In this setting, thrusts preferentially develop along the indenter front where they are rectilinear in plan view, and merge with strike-slip faults at both indenter lateral margins. This case does not fit with our field data. The case of symmetric oblique indentation shown in Fig. 14B corresponds to a block having the two sides oriented in the same way (45° in this example) as its motion vector. A symmetric arcuate belt develops in front of the corner and at both sides, with a possible minor component of transcurrence along the sides. This case does not enable explaining the differences along arc-strike of the tectonic regime in our study area. Finally, the case of asymmetric oblique indentation shown in Fig. 14C and D, refers to a movement direction of the indenter block at 30° with respect to the normal to indenter front. In plan view, the thrust traces are curved around the indenter corner and tend to be parallel to the indenter margins at the sides. Along the frontal margin of the indenter (left hand side in Fig. 14C and D), a wider belt of folds and thrusts develops, whereas in front of the lateral indenter margin (right side in Fig. 14C and D), a narrow belt of en-échelon structures develops. This difference is due to the lateral indenter margin forming a much greater angle (60°) from movement direction compared with front margin (30°). In this case, orogen-parallel extension does occur. All our constraints derived from field data match the model presented in Fig. 14D. We observed: (i) arc-parallel extension distributed along the arc; (ii) the lack of a long rectilinear front; (iii) a wider distribution of folding along the western side; (iv) the presence of folds

with a geometry more compatible with pure shear along the western margin; (v) and a fold pattern fully compatible with a left-lateral simple shear along the eastern margin. We thus suggest that the indenter is given by an overthrusting upper crustal block with a triangular shape in plan view. It is characterised by an angle between its main motion vector and the block margins, which is greater at the eastern side than at the opposite western side. The eastern margin might correspond to a lateral ramp, whereas the main frontal ramp should be located in proximity of the Tsaishi fold and further westward. These geometries allow us to speculate that the main motion vector of this indenter should be to the SSW. As previously mentioned, this overthrusting crustal block is limited downward by ramp-and-flat surfaces that should coincide with the decollement planes recognized in various crustal sections by Banks et al. (1997) and Tibaldi et al. (2017a, b). Based on these authors, the main and deeper decollement plane is located at a depth of 3–5 km below the Rioni fold-and-thrust belt, and tends to become deeper northwards, below the Greater Caucasus, reaching a depth of about 10 km. Our results highlight the general importance of the interaction between the indenter’s movement direction and the orientation of the indenter margins, with respect to other possible variables. In particular, an oblique asymmetric indentation is coherent with our field data and also with a much similar setting, observed in the southern Eastern Carpathians arcuate fold-and-thrust belt in Romania (Zweigel et al., 1998). A similar scenario has been recognized also for some phases of development of the external arc of the western Alps in France-Switzerland (Lickorish et al., 2002), or it might be envisaged for the northwestern Taiwan fold-and-thrust belt (Lacombe et al., 2003). 5. Conclusions We performed detailed field surveys in the Rioni Basin uplifted area, at the southern foothill of the western Greater Caucasus, in the Republic of Georgia. Here, marine and continental deposits of JurassicCretaceous-Neogene age were locally folded and uplifted, giving rise to 242

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an arcuate fold-and-thrust system with upper crustal rocks transported towards the foreland, like it happens in classical salients. Folds along the front are south-vergent and trend NW-SE, E-W and NE-SW. In the western part of the study area, folds are sinuous in plan view, whereas to the east they show a left-stepping, en-échelon geometry. In the interior of the uplifted area, folds are symmetric, trend NW-SE and NESW, and are more widespread to the W. All frontal folds are affected by dominant strike-slip faults and only very subordinate reverse faults. Microtectonic data were collected at 248 faults, and paleostress tensors were calculated for the various phases of deformation. The greatest principal stress σ1 is mostly horizontal, and trends about WNW-ESE (parallel to the Greater Caucasus trend), NE-SW, and about N-S. The intermediate principal stress σ2 is mostly vertical, consistent with the dominance of transcurrent faults; right-lateral strike-slip faults have various orientations, but they cluster in the NW-SE and NE-SW trends. Left-lateral strike-slip faults mostly strike NW-SE, NE-SW and NNE-SSW. At a few sites, the σ2 is horizontal: reverse faults strike along two main trends, NE-SW and WNW-ESE. Crosscutting relationships between faults and other criteria allowed to identify four main deformation phases. Phase 1 took place during the late Eocene, based on regional correlations, and was characterised by σ1 trajectories between NNW-SSE and NNE-SSW. Phase 2 shows a strong rotation of the σ1 axis that attained a WNW-ESE to E-W orientation. This phase took place at the beginning of the Miocene. Both these phases predate folding in the studied area. Phase 3 was characterised by NNE-SSW to NE-SW σ1 trajectories; it postdates folding that, in the study area, started through preliminary, minor deformation during the Oligocene and fully developed since mid-Miocene times. Phase 4 is marked by parallel σ1 trajectories trending from N-S to NNE-SSW; it is still active, as testified to by the occurrence of earthquakes in the study area and the presence of major structures, both folds and faults, bearing evidence of Quaternary deformation. The arcuate belt shows along-strike variations in tectonic regime and deformation geometries, plus belt-parallel stretching. Through a comparison of all field data with published analogue models, we propose a geometric explanation for the formation of the arcuate belt based on oblique asymmetric indentation of an upper crustal blocks that moved dominantly to the SSW.

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