Introduction

The Kyrenia Range, Cyprus in the easternmost Mediterranean region (Fig. 1a) has the potential to document the tectonic development of the Late Mesozoic-Pleistocene northerly, active continental margin of the Southern Neotethys (Cleintuar et al. 1977; Robertson and Woodcock 1986; Harrison et al. 2004a, 2012; Harrison 2008; Robertson et al. 2012a, b, 2014; McCay and Robertson 2013; Palamakumbura et al. 2015). The Kyrenia Range is located in a region of incipient continental collision (Kempler and Garfunkel 1994; Robertson 1998a, b; Vidal et al. 2000; Jolivet and Faccenna 2000; Harrison et al. 2004a), such that its structural development can be unravelled in exceptional detail, with implications for other parts of the Eastern Mediterranean and elsewhere. In this paper, we present the first detailed micro-structural analysis and interpretation of the key central segment of the Kyrenia Range.

Fig. 1
figure 1

a Tectonic setting of the eastern Mediterranean. Closed arrowheads indicate areas of crustal contraction. Open arrowheads indicate subduction zones. Large arrows indicate the plate motions of Africa and Arabia relative to a fixed Eurasia (McClusky et al. 2000). The Eratosthenes Seamount is shown to the south of Cyprus (in light grey). AM Anaximander Mountains, CB Cilicia Basin, DSFZ Dead Sea Transform Fault Zone, EAFZ, East Anatolian Fault Zone, FR Florence Rise, HB Herodotus Basin, IA Isparta Angle, IB Iskenderun Bay, LB Levant Basin, LLT Latakia-Larnaka-Tartus ridge, NAFZ North Anatolian Fault Zone, PT Pliny Trench, ST Strabo Trench. b Outline geology of the central Kyrenia Range, modified after Baroz (1979). The classification of the Quaternary units is after Palamakumbura (2015) and Palamakumbura et al. (2015)

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Structural and stratigraphical setting

The Kyrenia Range is dominated by four main sedimentary successions of Mesozoic–Cenozoic age. These are separated by unconformities which are dated as intra-Late Cretaceous, intra-Mid-Eocene and the intra-Messinian–Early Pliocene (Fig. 2). The successions are: (1) Triassic to Cretaceous shallow-water platform carbonates that were metamorphosed under greenschist facies conditions during the Late Cretaceous; (2) Late Cretaceous to Middle Eocene pelagic and re-deposited limestones, basic to silicic volcanogenic rocks and sedimentary melange (~olistostromes); (3) Late Eocene to Late Miocene sequences of non-marine to relatively deep-marine siliciclastic sandstones and mudrocks; and (4) Pliocene shelf-depth marls, mudrocks and minor rudaceous rocks. The four sequences are overlain by Quaternary, non-marine to coastal sediments (Palamakumbura 2015). All of the four stratigraphic successions include structures that shed light on the tectonic development of the central Kyrenia Range.

Fig. 2
figure 2

Stratigraphy of the Kyrenia Range (modified after Robertson et al. 2012b). The traditional stratigraphic names are given first, with their Turkish language equivalents in parentheses. Note the presence of major unconformities that resulted from regional deformation events. Here, and elsewhere, the time scale of Gradstein et al. (2012) is used

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The Kyrenia Range is traditionally divided into four segments: the western range, the central range, the eastern range and the Karpas Peninsula. These have many common geological features, but also some differences. The central Range (Fig. 1b) was chosen for study because of its central location and because it exposes the Mesozoic ‘core’ of the range, which is absent or poorly developed elsewhere. Structural data were mainly collected along the crest and northern flanks of the central range. The structure of the southern front of the Kyrenia Range has been discussed elsewhere (McCay and Robertson 2013). All of the available structural evidence is taken into account in the interpretation and discussion of our results in the regional setting.

Methodology

Three related types of evidence are used here. First, N–S structural traverses were made with the aim of understanding the large-scale structure in relation to the known stratigraphic development. For example, can any of the large-scale structural features (e.g. thrusts) be related to the four main stages of stratigraphic development separated by the unconformities? Knowledge of the stratigraphy allows the structural development to be ‘backstripped’ to reveal the timing of formation of specific structural features. Secondly, we have collected and analysed a large data set of small-scale structures including faults, folds and other structures from which stress regimes can be inferred. The third type of evidence comes from offshore and onshore seismic reflection data, other geophysical data, boreholes and regional comparisons.

Structure of the central range

Large-scale structure

Baroz (1979) provided 42 ~N–S sections of the Kyrenia Range, of which 21 are for the central range. These sections indicate a dominantly south-vergent, imbricate structure, which features in all subsequent structural interpretations (Robertson and Woodcock 1986; Harrison et al. 2004a; Calon et al. 2005a, b; McCay 2011; Robertson et al. 2012b, 2014). Baroz (1979) inferred numerous imbricate thrust slices of variable scale and geometry. However, a more recent interpretation of the seismic profiles that are located offshore to the west of the Kyrenia Range (Calon et al. 2005a) suggest that the internal structure of the Kyrenia Range is relatively more organised and less complex than envisaged by Baroz (1979). Baroz (1979) envisaged complex sub-surface geometries, which are not visible on the offshore seismic profiles. However, the resolution of the seismic profiles is inadequate to reveal outcrop-scale structure and further work is needed to test Baroz’s (1979) interpretations.

Five reconnaissance cross sections were made of the central range, making use of stratigraphic information, including way-up evidence, and micro- to meso-scale structural evidence (Fig. 3). Additional cross sections of the eastern range are given by Robertson et al. (2014). On a local scale, the structure of the central range appears to be complex with numerous thrusts, reverse faults, normal faults and strike-slip faults, many of which may have been reactivated during successive deformation phases. However, the large-scale structure becomes apparent when the tectonostratigraphy is taken into account (Fig. 2).

Fig. 3
figure 3

Sketch cross sections across the central Kyrenia Range (see Fig. 1 for locations). Modified from Robertson et al. (2014). See text for explanation

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The main features of the traverses are as follows.

Section (a), Selvili Tepe (Kornos) area

There is a clearly defined northward-dipping imbricate structure (Fig. 3a) involving repeated slices of Mesozoic carbonate rocks, latest Cretaceous-Eocene carbonates and volcanics, and Late Eocene-Late Miocene marine siliciclastic sediments. The Late Eocene-Late Miocene sedimentary rocks are restricted to near the base of the thrust stack in the east. The topographically highest thrust sheet, near the crest of the range, is deformed into a large-scale south-verging recumbent fold. Thrust sheets of Mesozoic carbonate rocks, latest Cretaceous-Eocene pelagic limestones and Late Eocene-Late Miocene ages are deformed into north-verging back-thrusts. As a result, along this part of the northern flank of the range, the Late Eocene-Miocene siliciclastic sedimentary rocks are overturned.

Section (b), Lapta area

This shows a similar imbricate structure (Fig. 3b). However, here, the Late Eocene-Late Miocene siliciclastic sedimentary rocks are not exposed at the base of the thrust stack. Late Eocene-Miocene siliciclastic sedimentary rocks along the north margin of the range are steeply northward dipping but not overturned in this area.

Section (c), Dikmen (Ano Dikmen) to Beylerbey (Bellapais) area

This section shows the typical northward-dipping geometry of the thrust sheets (Fig. 3c). Also, the uppermost thrust sheets, including the unconformity between the latest Cretaceous-Middle Eocene and the Late Eocene-Miocene successions, are overturned, similar to Section a. However, the lower part of the thrust stack is dominated by a large sheet of Mesozoic meta-carbonate rocks which, unusually, dips southwards at a relatively shallow angle. The bedding in this sheet is truncated by thrusts above and below. This structure hints at a complex multi-phase development involving re-imbrication and steepening.

Section (d), Beşparmak (Pentadactylos) area

The central range is relatively wide in this area (Fig. 3d). A substantial thrust sheet of Late Cretaceous-Eocene pelagic carbonate, mafic igneous rocks and sedimentary melange in the south is deformed into a large-scale south-verging fold structure, which is internally faulted. The mountain to the north (Beşparmak (Pentadactylos)) is made up of northward-dipping Mesozoic carbonate rocks.

Section (e), West of Antifonitis monastery

This transverse is located close to the junction with the eastern range (Fig. 3e). The Mesozoic-Late Eocene succession is extensively overturned along the northern flank of the range in this area. The eastern range shows evidence of a contrasting structural style which reflects the absence of structurally competent thick Mesozoic platform carbonates but the addition of large volumes of incompetent sedimentary melange (‘olistostromes’; McCay and Robertson 2013; Robertson et al. 2014).

When the stratigraphy is taken into account, it becomes clear that a relatively small number of large imbricate slices are present, comparable in scale with those inferred from offshore seismic data (Calon et al. 2005a). Most of the major thrusts are sub-parallel to the large-scale stratification of the Trypa (Tripa) Group meta-carbonate rocks, yet locally the internal stratigraphy of the meta-carbonates is transected by thrusts at various angles (e.g. Fig. 3d). The major thrusts typically emplace Triassic-Cretaceous meta-carbonate rocks of the Trypa (Tripa) Group over younger sequences, namely the Maastrichtian-Early Eocene pelagic carbonates and volcanics of the Lapithos (Lapta) Group and/or the Middle Eocene debris-flow units of the Kalograia-Ardana (Baçeli-Ardahan) Formation (e.g. Fig. 4a). Major thrusts generally cut stratigraphically upwards towards the south.

Fig. 4
figure 4

Field photographs of large-scale structural features within the central Kyrenia Range. a Highly sheared pelagic limestones and basaltic volcanics of the Late Cretaceous Malounda (Malıdağ) Formation, thrust southwards over turbidites and debris-flow deposits of the Early Eocene Kalograia-Ardana (Bahçeli-Ardahan) Formation; ~1.2 W of Kythrea (Değirmenlik) spring. b NW–SE-trending horst block of meta-carbonates of the Mesozoic Hilarion (Hilaryon) Formation, overlain by the Maastrichtian Malounda (Malıdağ) Formation, ~2 km S of İlgaz (Fteriha). c Transverse fault scarp terminating the central range, west of Selvili Tepe (Kornos). d ~N–S-trending steep fault juxtaposing meta-carbonates of the Hilarion Formation (right) with highly sheared pelagic carbonates of the Maastrichtian Malounda (Malıdağ) Formation; photograph taken at an oblique view to Fig. 5c. The top of the prominent fault cuts the topography in the position of the letter c. e ~NNW–SSE-trending eroded fault scarp in meta-carbonate rock of the Hilarion (Hilaryon) Formation (background); W side of road across the central range, ~1.5 km NW of Beşparmak (Pentadactylos). f ~N–S-trending steep faults cutting meta-carbonates of the Hilarion Formation; crest road SE of Beşparmak (Pentadactylos), towards the Antifonitos monastery

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The Mesozoic to Miocene strata of the Central Range are tilted at variable angles, generally towards the north. The Mesozoic-Paleogene units typically dip northwards at ~60°. The rugged topography has generally resulted from the erosion of strongly deformed and jointed, lithologically competent carbonate rocks (e.g. Fig. 4b, c). The primary bedding, which is commonly obscure, dips northwards at generally moderate to steep angles. The Mesozoic-Paleogene units and also the basal part of the unconformably overlying Kythrea (Değirmenlik) Group are steeply dipping to overturned in some places along the northern flank of the central range. The strata are imbricated on a large scale and also commonly folded about gently plunging axes. The folds along the southern flank of the central range (Fig. 5a) are generally south-vergent (McCay and Robertson 2013), although there is evidence of north-verging folding along the northern flank of the range in several areas.

Fig. 5
figure 5

Field photographs of typical outcrop-scale structures within the Triassic-Late Miocene lithologies of the central Kyrenia Range. a Brecciated carbonate rocks of the Trypa (Tripa) Group, thrust southwards over Eocene debris flows of the Kalograi-Ardana (Bahçeli-Ardahan) Formation. South of the main road, c. 1 km west of Değirmenlik (Kythrea). b ~NNW–SSE-trending fault scarp in meta-carbonate rock of the Hilarion (Hileryon) Formation. Most of the fault scarp is eroded; however, the recently exposed lower several metres above the ground surface shows dip-slip slickenlines; ~0.5 km NW of Beşparmak (Pentadactylos); N flank of central range (the triangular-shaped ridge is ~0.5 km across at the base). c Typical strike-slip slickenlines with left-lateral offset that characterise numerous ~N–S faults cutting the meta-carbonates of the Hilarion (Hileryon) Formation; crest road east of Hilarion castle. d Typical oblique-slip, strike-slip slickenlines, in this case left-lateral, cutting the meta-carbonates of the Hilarion (Hileryon) Formation. Oblique-slip slickenlines are associated with many of the ~N–S high-angle faults; 1.5 km SE of Karaman (Karmi); N flank of central range. e Rarely observed two sets of slickenlines on a NW–SE-trending fault plane (oblique-slip and normal); cutting pink pelagic carbonates of the Maastrichtian Malounda (Malıdağ) Formation; ~1 km S of Antifonitis monastery. f ~NE–SW tending, gently plunging fold, cut by sub-parallel steeply inclined fault; within the Palaeocene–Eocene Ayios Nikolaos (Yamaçköy) Formation; 3 km east of Alsancak (Karavas), N flank of central range; g ~N–S high-angle fault cutting the basal layers of the Bellapais (Beylerbey) Formation; E side of the Nicosia (Lefkoşa) to Girne (Kyrenia) highway; ~4 km S of central Girne (Kyrenia). h ~N–S-trending joint planes with minor displacement, cutting rare sandstones of the Mid-Miocene Trapeza (Esentepe) Formation; ~1 km S of Alagadi beach (~Platimatis valley); N flank of central range; i ~N–S-trending high-angle faults cutting Late Miocene sandstones of the Davlos (Kaplıca) Formation; roadside above Esentepe (Agios Amvrosios); N flank of central range; j ~E–W-trending zone of high-angle shearing and folding within the Late Eocene-Early Miocene sandstones and mudrocks of the Bellapais (Beylerbey) Formation, near Yeşiltepe (Elia Kerinias); N flank of central range; k Steeply plunging, ~NE–SW-trending chevron-type fold in mudrocks and sandstones of the Bellapais (Beylerbey) Formation; several kilometres east of Aslancak (Karavas); l W-verging upright fold with ~N–S-trending fold axis; within sandstones and mudrocks of the Bellapais (Beylerbey) Formation; associated with sub-parallel zone of shearing; road over crest of range ~2.5 km W of Beşparmak (Pentadactylos)

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The contact between the Triassic-Cretaceous meta-carbonate rocks (and minor interbedded lithologies) of the Trypa (Tripa) Group and the overlying unmetamorphosed lithologies of the Lapithos (Lapta) Group is a north-dipping unconformity. This is commonly exploited by reverse faulting, thrusting and occasionally by normal or strike-slip faulting. The unconformity is well exposed within the major thrust sheet that forms the backbone of the central range (e.g. Fig. 3a). This contact can be traced for ~8 km along the upper northern flank of the central range in the west (Robertson et al. 2012b).

The contact between the lithologies of Mid-Eocene and Late Eocene age is an unconformity, which is locally exposed along the northern and southern flanks of the central range. Along the northern flank, the unconformity ranges from northward dipping, to steep, to locally southward dipping as a result of northward back-thrusting (Fig. 3c). Along the southern flank of the central range, this unconformity is largely eroded or cut out by thrusts, although it remains well exposed in the eastern range (Baroz 1979; McCay and Robertson 2013).

Geological maps of the central range (Ducloz 1972; Baroz 1979; Hakyemez et al. 2000) reveal numerous transverse structures at a variety of scales. These range from ~N–S to NE–SW in the west and ~NW–SE in the east. In the west, the central range as a whole is transected and offset southwards by an important ~N–S to NNE–SSW-trending fault lineament (Fig. 4c, d; see Fig. 1 for location). To the east, the central range is effectively terminated by another major ~N–S transverse fault lineament (Fig. 1). These two major lineaments effectively delimit the central range. In between, large transverse lineaments have been exploited by erosion to form topographical depressions. These depressions include a prominent ENE–WSW-trending depression (valley and pass) south of Kyrenia (Girne), which effectively splits the range into two segments and a significant NNW–ESE-trending lineament further east, in the Beşparmak (Pentadactylos) area (Figs. 4e, 5b; see Fig. 1 for locations). Numerous smaller-scale lineaments with similar trends also transect the central range (e.g. Fig. 4b). The timing of the development of the transverse structures is debateable: a relatively young, Late Miocene to Plio-Quaternary age was inferred by Harrison et al. (2004a), mainly because similar structures were mapped by these authors within the Miocene sedimentary rocks of the Kythrea (Diğirmenlik) Group in the Mesaoria (Mesorya) Basin. Alternatively, a pre-Late Eocene age was suggested by McCay and Robertson (2013), mainly because some of the transverse structures did not appear to cut the Miocene sedimentary rocks (see Baroz 1979). However, fault reactivation was recognised to have played an important role. The two alternatives were tested during this work (see below).

The combined information indicates that the large-scale structural development of the central range was dominated by three phases of convergence: (1) Late Cretaceous deformation associated with greenschist facies metamorphism (semi-ductile to brittle); (2) Mid-Eocene thrusting and folding (semi-ductile); and (3) Late Miocene-Early Pliocene thrusting and folding (semi-brittle).

Outcrop-scale structure

Mesozoic succession

The Triassic-Cretaceous lithologies are exposed within the thickest thrust sheets that dominate the crest of the central range (Fig. 1). The Late Triassic to Early Cretaceous succession (Saint Hilarion (Hileryon) Formation) is composed of medium- to thick-bedded, re-crystallised shallow-water carbonate rocks, ranging from finely crystalline to coarsely crystalline (i.e. marble). These rocks are structurally competent, extensively faulted but rarely show evidence of folding on an outcrop scale. Parts of the succession, notably the Triassic Sikhari (Kaynakköy) Formation, are extensively dolomitised (Ducloz 1972; Baroz 1979; Robertson and Woodcock 1986). Where well exposed, for example, along the road across the central range west of Beşparmak (Pentadactylos) (Fig. 1), the Mesozoic stratigraphy is relatively intact and does not show obvious evidence of internal tectonic disruption on an outcrop scale (Fig. 6).

Fig. 6
figure 6

True-scale section across part of the Central Kyrenia Range ~2 km west of Hilarion (Hileryon) castle (see Fig. 1 for line of section). This section is relatively simple with two thrust sheets of Mesozoic meta-carbonate rocks, overlain by Maastrichtian-Eocene pelagic carbonates and volcanics of the Trypa (Tripa) Group

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Volumetrically subordinate, dark coloured, locally organic-rich shales (mudrocks) are exposed locally, especially in the Triassic succession. The incompetent sedimentary rocks typically exhibit a bedding-parallel cleavage and commonly acted as local zones of shearing and folding.

After deposition, the entire platform carbonate succession was metamorphosed to greenschist facies during the Late Cretaceous, as indicated by the mineral assemblages in the interbedded pelitic sedimentary rocks (Baroz 1979; Robertson et al. 2012b). The schistosity in the mudrocks is assumed to relate to this event.

The Mesozoic meta-carbonate rocks are cut by an array of generally steeply dipping joints, faults and shear planes, many of which are orientated ~N–S (Fig. 5b). These faults are assumed to post-date metamorphism. Where present, kinematic evidence in the form of slickensided fault surfaces (with exposed slickenlines) indicates a strong strike-slip to oblique strike-slip component of displacement (Fig. 5c, d).

A characteristic feature of the platform carbonates is that they are strongly brecciated in many outcrops (Robertson and Woodcock 1986; Robertson et al. 2012b). The brecciation ranges from extreme, such that no trace of the original bedding can be discerned, to moderate where bedding is still visible and to minor where primary bedding is little affected. The timing of tectonic brecciation is known to be Late Cretaceous (Robertson and Woodcock 1986), because the brecciation took place after the youngest dated deposition of the shelf carbonates (Cretaceous) and the Late Cretaceous (Maastrichtian) age of the unconformably overlying cover succession. Clasts at the base of the Maastrichtian cover succession have been reworked from underlying previously brecciated meta-carbonates (Robertson et al. 2012b). The tectonic breccias have previously been explained, either as a response to strike-slip faulting (Robertson and Woodcock 1986) or as a response to extensional exhumation (Robertson et al. 2012b). However, both processes may not be mutually exclusive; i.e. exhumation could have taken place in a transtensional setting.

Latest Cretaceous-Middle Eocene succession

The succession overlying the metamorphosed Mesozoic platform carbonates begins with transgressive, relatively deep-marine pelagic carbonates (Baroz 1979; Robertson et al. 2012b). The lower part of this succession includes localised intercalations of sandstone–siltstone turbidites (Ducloz 1972; Baroz 1979; Robertson and Woodcock 1986). The terrigenous sediments were recently re-defined as the Kiparisso Vouno Member of the Maastrichtian Malounda (Malıdağ) Formation (Robertson et al. 2012b). These relatively thin (tens of metres), structurally incompetent sedimentary rocks are commonly exploited by thrusting. Higher in the succession, debris-flow deposits (up to several metres thick) are dominated by clasts derived from the underlying meta-platform succession. Numerous intercalations of alkaline intra-plate-type basaltic lavas occur throughout the Maastrichtian-aged succession (Robertson and Woodcock 1986; Huang et al. 2007; Huang 2008). These volcanic rocks are structurally competent compared to the interbedded pelagic carbonates and are often bounded by shear zones and cut by faults. The dominant trend of this faulting is NW–SE (with subordinate ENE–WSW trends), indicative of overall N–S compression. There are also congruent sinistral and dextral NE–SW-orientated faults, and sinistral NW–SE-orientated faults, as discussed further below.

The mixed pelagic carbonate-limestone debris-flow volcanic succession passes upwards into generally more argillaceous, thinner-bedded chalks, marls and interbedded volcanics of Palaeocene-Early Eocene age, represented by the Ayios Nikolaos (Yamaçköy) Formation. The formation includes additional intra-plate-type basaltic rocks (Robertson and Woodcock 1986; Huang et al. 2007; Huang 2008). In the central range, the Ayios Nikolaos (Yamaçköy) Formation then passes depositionally upwards into marls, calciturbidites and debris-flow deposits of the Middle Eocene Kalograia-Ardana (Bahçeli-Ardahan) Formation (Baroz 1979; Robertson et al. 2014). The clasts in the gravity flows are similar to the lithologies of the underlying Mesozoic-Paleogene succession. However, other clasts include boninite-type ophiolitic rocks and also shelf-facies limestones with large foraminifera (i.e. Nummulitic limestones). These additional rock types appear to have been derived from north of the Kyrenia Range (Robertson et al. 2014).

The Maastrichtian-Eocene succession is cut by an array of mostly high-angle faults of variable geometry, ranging from dominantly NNE–SSW to ENE–WSW and NW–SE orientations. Kinematic evidence in the form of slickenlines on slickensided fault surfaces indicates a variety of displacements, ranging from reverse (oblique-slip; Fig. 5e), to normal, to pure strike-slip. Two sets of slickenlines are rarely preserved on a single fault plane (Fig. 5e), providing localised evidence of fault re-activation. In addition, a pressure solution cleavage, ranging from bedding-parallel to bedding-oblique, is well developed in the debris-flow units, in which most of the clasts are composed of competent limestone or dolomite. The strike of the cleavage mainly strikes ENE–WSW, whereas a subordinate, congruent data set strikes NNW–SSE.

In places, the marls and chalks are sheared to the extent that primary bedding is destroyed. The pelagic chalks are folded and faulted, whereas less competent marls and clayey limestones preferentially exhibit shearing and ductile deformation. Good examples of highly sheared chalk–marl sequences occur along the southern margin of the central range, near Kythrea (Değirmenlik) and along the northern margin of the central range, near Karşıyaka (Vassilikos).

The Maastrichtian to Mid-Eocene succession is locally folded, mostly on a tens-of-centimetre to several-metre scale. Such folds are abundant along both the southern and the northern flank of the central range, especially near the top of the Maastrichtian-Middle Eocene succession. The folds are characterised by a range of ~NW–SE to E–W orientated axial planes and are both westerly and northeasterly vergent. Some of these folds are cut by high-angle faults (Fig. 5f).

Late Eocene-Late Miocene succession

The stratigraphy of the Late Eocene-Late Miocene succession is cut by a variety of structures, mostly high-angle faults of several different orientations. However, the different formations that crop out at different distances from the core of the Kyrenia Range are affected to variable extents (as seen along the northern flank of the central range).

The Late Eocene-Late Miocene Kythrea (Değirmenlik) Group begins, above the basal unconformity, with non-marine conglomerates up to several tens of metres thick (Baroz 1979; Robertson and Woodcock 1986; McCay and Robertson 2012). This relatively competent unit was derived from lithologies similar to those exposed beneath, together with numerous diabase/microgabbro clasts of boninitic composition. These intrusive igneous rocks were interpreted to have been derived from a supra-subduction zone-type ophiolite (McCay 2011) that was located to the north of the Kyrenia Range (Robertson et al. 2014).

The basal coarse-grained clastic sedimentary rocks pass upwards into less competent Oligocene-aged mudrocks, sandstones and debris-flow conglomerates (Bellapais (Beylerbey) Formation) and then into sandstone turbidites, mudrocks and hemipelagic marls of the Klepini (Arapköy), Flamoudi (Tirmen) and Panagra (Geçitköy) Formations. Above this come generally thinner-bedded, finer-grained mudrocks, sandstone turbidites, marls and sapropels (see Talyforth et al. 2014) of Mid-Late Miocene age (Trapeza (Esentepe) Formation). The above five formations show similar styles of deformation. Much thicker-bedded, relatively competent sandstones (gravity flows) and incompetent mudrocks characterise the Late Miocene (Davlos (Kaplıca) Formation). No higher units of the Kythrea (Değirmenlik) Group are exposed along the northern flank of the central range.

In contrast, the succession continues into the Messinian, including evaporites, in places along the southern flank of the range, although not in the area of the central range studied here. In general, the Late Eocene-Miocene succession along the southern flank of the central range is highly deformed into small-scale imbricate thrust sheets, associated with folds, which mostly verge southwards. The lowest exposed, major thrust is represented by the Kythrea (Değirmenlik) Fault (McCay and Robertson 2013; see below). The succession in the south is cut by an array of faults, which prominently trend ~N–S. These faults are interpreted to indicate dominantly southward displacement in a left-lateral (transpressional) setting, coupled with extensive fault reactivation (McCay and Robertson 2013). In addition, large-scale south-vergent nappe structures fold the Late Eocene-Miocene rocks and older units within the western range outside the study area (Baroz 1979; Robertson et al. 2012a; McCay et al. 2013).

Extensive new structural data were obtained from the Late Eocene-Late Miocene succession exposed along the northern flank of the central range. Faults are preferentially exposed in relatively lithified intervals, with the deformation generally decreasing up-section. The exposures are in mostly steam valleys and locally in areas of urban development, for example west of Kyrenia (Girne), near İlgaz (Fteriha) and also east of Kyrenia near Esentepe (Agios Amvrosios). The succession is relatively steeply dipping to locally overturned in the south (near the basal unconformity) but becomes generally less steeply dipping stratigraphically upwards towards the coast (e.g. west of Kyrenia near Karaman (Karmi)). In places, both the unconformity and the overlying basal clastic sedimentary rocks are sub-vertical to slightly inverted, e.g. near Beylerbey (Bellapais).

The Late Eocene to Late Miocene succession is cut by an array of faults, which trend predominantly NNW–SSE to NE–SW [a congruent set of faults trends WSW–ENE to NW–SE (e.g. Fig. 5g)]. A zone of intense approximately E–W shearing and associated folding was observed within the Bellapais (Beylerbey) Formation west of Kyrenia (near İlgaz (Fteriha); see Fig. 5j). The overlying formations are characterised by less intense brittle structures, trending predominantly NW–SE, with a subordinate congruent set trending NE–SW. Additional faults trend NNW–SSE. The mudrock-dominated formations (e.g. Trapeza (Esentepe) Formation) are characterised by joints with centimetre-scale offsets that are observable within rare relatively competent sandstones (Fig. 5h). The well-cemented sandstones of the overlying Davlos (Kaplıca) Formation exhibit micro- and meso-scale high-angle faults (Fig. 5i), which are more or less invisible within the interbedded poorly lithified mudrocks. Internally, massive sandstones are characterised by networks of granulation seams. These are indicative of deformation while the sediments were still water-rich and poorly consolidated. Folds in the Late Eocene-Miocene succession as a whole are rare and localised (e.g. Figs. lb, 5k).

The faults cutting the succession exposed along the northern flank of the central range are less pervasive and tend to be localised, compared to those documented from the southern flank of the central range (McCay and Robertson 2013). However, rare ~N–S-trending faults are present in all of the locations studied. Locally, ~N–S faults are observed to cut folds within the Paleogene pelagic carbonates (e.g. 3 km east of Alsancak (Karavas); see Fig. 5f). Also, the unconformity between the Paleogene, mostly carbonate sedimentary rocks, and the overlying Late Eocene–Oligocene siliciclastic sedimentary rocks is cut by ~N–S high-angle faults (e.g. near Beylerbey (Bellapais)). The faulting within the Late Eocene-Late Miocene sedimentary rocks is most intense in outcrops located near the generally northward prolongations of the major transverse lineaments that cut the central range. The most notable of these are the Boğaz lineament south of Kyrenia (Girne) and the Beşparmak (Pentadactylos) lineament further east (see Fig. 5b). A number of the smaller-scale transverse lineaments also extend at least into the Bellapais (Beylerbey) Formation, as indicated by locally intensified faulting (e.g. at Lapta (Lapithos) spring).

In summary, some thrusts and folds affect the Kythrea (Diğirmenlik) Group and, therefore, must relate to the Late Miocene-earliest Pliocene phase of compressional deformation. Thrusts of a similar trend also cut the subjacent Paleogene and Mesozoic sedimentary rocks and so at least some of these structures are likely to be of Late Miocene-earliest Pliocene age. However, these structures could also reflect the re-activation of older structures, of Mid-Eocene and possibly Late Cretaceous age. For example, west of Lapta (Lapithos), Mesozoic meta-carbonates are back-thrust northwards over a thin slice of Palaeogene carbonates and marls of the Ayios Nikolaos (Yamaçköy) Formation (see Figs. 7a–f, 8). These sedimentary rocks are, in turn, back-thrust over siliciclastic sedimentary rocks of the Late Eocene–Oligocene Bellapais (Beylerbey) Formation (Fig. 8). The available kinematic evidence, as observed along the two major back-thrusts in this area, indicates dominantly oblique northward vergence (Fig. 7a, b) with a strong left-lateral (top-to-the-west) component of displacement (Fig. 7b, d, f). The relatively competent carbonate rocks of the hanging-wall (re-crystallised limestone) are tectonically brecciated within tens of centimetres of the thrust plane (Fig. 7c). Small-scale ~N–S faults transect and slightly offset the thrust plane (Fig. 7e). The available evidence, therefore, points to oblique sinistral displacement initially during the Mid-Eocene deformation phase, followed by fault reactivation, northward back-thrusting and ~N–S high-angle faulting during the Late Miocene-earliest Pliocene deformation phase.

Fig. 7
figure 7

Field photographs illustrating the important role of northward back-thrusting coupled with westward displacement (sinistral transpression) of inferred Late Miocene-earliest Miocene timing. All photographs are views generally to the south. 1 km WSW of Lapta (Lapithos) (GPS 36S 14295 UTM 39 09955). a Steeply S-dipping slickensided fault scarp cutting meta-carbonate rocks of the Hilarion (Hileryon) Formation; thrust obliquely northwards over highly sheared clayey pelagic limestones of the Palaeocene–Eocene Ayios Nikolaos (Yamaçköy) Formation. b As above; note the westward entrainment of the footwall clayey pelagic limestones. c Tectonic brecciation of the relatively competent hanging-wall meta-carbonate rocks. d Intense shearing of the relatively incompetent footwall wall clayey pelagic carbonates; e ~N–S steep faults in hanging-wall meta-carbonate rocks, cutting through into sheared clayey pelagic limestone in the footwall beneath. f West-directed shear bands (C–S fabric) developed in clayey pelagic limestone of the footwall

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Fig. 8
figure 8

Local cross section of the northern flank of the central range, W of Lapta (Lapithos), showing the structural context of the features shown in Fig. 7. The Mesozoic meta-carbonates are folded and thrust northwards over a thin slice of Paleogene pinkish pelagic carbonates and marls. These sedimentary rocks are, in turn, thrust over a right-way-up sequence of Late Eocene–Oligocene siliciclastic sedimentary rocks. The cross section, the Angelier plots and the Dihedra plots and the associated field photographs (Fig. 7) illustrate the important role of back-thrusting of inferred Late Miocene-earliest Pliocene timing. Some of these structures appear to re-activate Mid-Eocene convergent structures

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Analysis of small-scale structures

Statistical analysis of small-scale structures (e.g. faults and folds) can be used to reconstruct palaeo-stress regimes. McCay and Robertson (2013) used this method to help constrain the deformation events affecting the Late Eocene-Late Miocene sedimentary rocks exposed along the southern flank of the Kyrenia Range. Similar outcrop-scale data were obtained in this study, mainly from the Mesozoic-Paleogene units exposed in the core of the central range. The data were mainly collected from road sections, particularly the mountain road that runs ~E–W along the crest of the central range. Additional data were collected from several roads running obliquely across the central range and also from Mesozoic to Miocene exposures along the northern flank of the central range.

In total, over 2885 structural measurements were made, of which 2688 are from faults and fractures, 116 from bedding planes, 79 from cleavage planes/C–S fabrics and 101 from folds (both fold axes and plunge). Slip data were obtained from 537 faults. By lithological unit, the fault data were further subdivided into the Mesozoic succession (Trypa (Tripa) Group) 1572 (360 with measurable kinematic indicators), the Maastrichtian-Middle Eocene succession (Lapithos (Lapta) Group), comprised of the Kalograia-Ardana (Bahçeli-Ardahan) 121 (38 with measurable kinematic indicators) and the Ayios Nikolaos (Yamaçköy) Formations 497 (87 with measurable kinematic indicators), and the Late Eocene-Late Miocene succession (Kythrea (Değirmenlik) Group) 435 (34 with measurable kinematic indicators). The principal stresses responsible for the generation and subsequent deformation of each fault set were calculated using the dihedral protocol, initially described by Angelier and Mechler (1977). The fault kinematics was determined using the program Tectonicsfp v.1.6.2 of Franz Reiter and Peter Acs. Each structural data point was assigned a co-ordinate, such that the entire data set could be interrogated spatially, which was necessary given the large amount of data collected.

The 2885 structural data set was treated as 140 localities, which were sub-divided into 16 subgroups. Eleven of the subgroups represent specific sections where large data sets were collected, whereas the remaining five subgroups represent specific trends, ages or units. By plotting data spatially, subsets of faults can be subdivided by orientation or age of the unit affected (i.e. ‘backstripped’) allowing similarities or differences within the data set to be evaluated, for example between individual large-scale thrust sheets and between different parts of the central range (Fig. 9). This allowed the structures affecting the different formations to be recognised and so ‘backstripped’ from the overall data population. In principle, structures that are common to several superimposed formations should post-date these particular stratigraphic units. However, lithological competency and fault re-activation can influence the nature of the structures present.

Fig. 9
figure 9

a Outline map of a selected (representative) area S of Beşparmak (Pentadactylos), showing the spatial distribution of the structural data collected in this region; b rose diagrams, Angelier plots, and Dihedra plots of the fault data sub-divided by locality; and c summary plots of the amalgamated fault data (as presented by region in Fig. 10)

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When the data set was analysed spatially and temporally several features emerged:

  1. 1.

    WNW–ESE- to ENE–WSW-trending normal faults affect the Mesozoic succession and to a lesser degree the Palaeogene and Neogene units.

  2. 2.

    NE–SW to ENE–WSW faults affect the Mesozoic, Paleogene and Miocene successions, as well exposed along the crest and the northern margin of the central range (Fig. 10). In addition, some faults of this trend also affect the Miocene sedimentary rocks exposed along the southern flank of the Central Range, in the Mesaoria (Mesarya) Basin (McCay and Robertson 2013). The stress regime that gave rise to this fault trend must, therefore, have existed as late as the Pliocene–Pleistocene, although it could also have existed earlier, with subsequent reactivation.

    Fig. 10
    figure 10

    Summary diagram of representative fault data from the central Kyrenia Range. The faults shown affect: a the Mesozoic Saint Hilarion (Hileryon) and Sykhari (Kaynnaköy) Formations (Trypa (Tripa) Group); b the Palaeogene Malounda (Malıdağ) and Ayios Nikalaos (A.N.) (Yamaçköy) Formations; c the Mid-Eocene Kalograia-Ardana (K-A) (Bahçeli-Ardahan) Formation (Lapithos (Lapta) Group); and d the Late Eocene-Late Miocene Bellapais (Beylerbey) and Davlos (Kaplıca) Formations (Mesaoria (Mesarya) Group). a Rose diagrams, illustrating the orientations of faults affecting the above units; b the resolved Dihedra plots of the fault data as sub-divided by locality. Note, the summary column includes data from outside the sub-groups shown to the left. All of the fault data are supplied in the supplementary data files. See text for ‘Discussion’

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  3. 3.

    NW–SE to NNW–SSE sinistral faults are present within the Mesozoic, Paleogene and Neogene units. The stress regime responsible for the generation of these faults may thus have existed as late as earliest Pliocene time based on the measurements made.

  4. 4.

    The faults affecting the Mesozoic-Neogene units show a gradual east–west change from ENE–WSW to WNW–ESE. This change in trend appears to reflect the arcuate shape of the range, consistent with relatively late-stage oroclinal bending. Such oroclinal bending could, in principle, have taken place during the Mid-Eocene and/or the Messinian-Early Pliocene phases of deformation. However, it could also have taken place progressively during Neogene to Recent time. Palaeomagnetic data show that the Karpas Peninsula in the far east of the island has undergone c. 30° anticlockwise tectonic rotation relative to the eastern range (Morris et al., this volume) (note: the central and western ranges did not yield interpretable palaeomagnetic data). The arcuate trend of the eastern part of the Kyrenia Range is, therefore, likely to relate to the inferred anticlockwise tectonic rotation.

  5. 5.

    The Paleogene units are affected by faults of many different orientations, with significant sub-populations of reverse, strike-slip and normal faults. The dominant trends are NNE–SSW, NW–SE and ENE–WSW. The NNW–SSE and NW–WE fault sets commonly exhibit slickensided surfaces, with slickenlines indicating an oblique sinistral component of movement along an ENE-WSW plane (σ1, 017/01; σ2, 285/53 σ1, 107/37; n = 171). The stress regime responsible for the generation of these faults could have originated during the Late Miocene-earliest Pliocene phase of deformation. However, given the stratigraphic relationships outlined above, it is likely to have already existed during the Mid-Eocene phase of deformation. Numerous folds affecting the Miocene sedimentary rocks along the southern flank of the Kyrenia Range reflect sinistral transpression during the Late Miocene-earliest Pliocene phase of deformation (McCay and Robertson 2013). However, there are also rare instances of this style of deformation persisting into the Pleistocene (McCay and Robertson 2013).

  6. 6.

    The deviatoric stress tensors that can be resolved for the compressional and transpressional structures that are inferred to have formed from the Eocene onwards are shown in Figs. 9b, c and 10b. To generate these stress tensors, the fault data were subdivided spatially, by geographic location, temporally, by age of affected unit and also kinematically, by the dominant type of displacement. To illustrate this approach, we show a selected data set from south of Beşparmak (Pentadactylos) (Fig. 9). The stress regime responsible for the generation of these faults, as indicated by the latest set of slickenlines on slickensided fault surfaces, is inferred to be N–S compression, which locally varied between NNW–SSE- and NNE–SSW-directed compression.

Regional-scale structural relationships

Additional structural information comes from south of the Kyrenia Range including outcrop and borehole data, and also from limited onshore and offshore seismic reflection data.

Evidence of structural shortening is found to the south of the Kyrenia Range, especially associated with the ~E–W-trending Kythrea (Değirmenlik) Fault (Ducloz 1972; Baroz 1979; Fig. 11). This fault places allochthonous rocks of the Kyrenia Range over parautochthonous Miocene siliciclastic sedimentary rocks within the Mesaoria (Mesorya) Basin. The Kythrea (Değirmenlik) Fault originated as a down-to-the-north convergent lineament that influenced sedimentation during Mid-Late Miocene time (Robertson and Woodcock 1986; McCay and Robertson 2013). Further south, the Ovgos (Dar Dere) Fault Zone (Harrison et al. 2004a; McCay and Robertson 2013; Fig. 11) exhibits additional important evidence of convergence and left-lateral strike-slip during different times (Harrison et al. 2004a; McCay and Robertson 2013; Fig. 11).

Fig. 11
figure 11

Regional N–S cross section based on Harrison et al. (2004a, b) and data sources cited by these authors. The section given by Harrison et al. (2004a, b) is extended southwards to the Troodos ophiolite outcrop using information from Kinnaird and Robertson (2012). The structure of the Kyrenia Range is based on Baroz (1979), Robertson et al. (2014) and this paper. The section is also extended northwards, offshore using data in Calon et al. (2005a). The stratigraphy differs north and south of the Ovgos (Dar Dere) Fault (McCay and Robertson 2012)

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Shallow and deep-level seismic reflection data from the westward, offshore extension of the Kyrenia Range are interpreted to indicate the presence of a series of northward-dipping thrust imbricates, which are folded in places. The deformed crust is overlain, unconformably by essentially undeformed sedimentary rocks of inferred Plio-Quaternary age (Calon et al. 2005a, b). The offshore data indicate that the imbricate structure persists to the south of the Kyrenia Range, including the westward prolongations of the Kythrea (Değirmenlik) Fault and the Ovgos (Dar Dere) Fault. A major down-to-the-north (extensional) high-angle fault is imaged to the north of the island, apparently accommodating part of the Plio-Quaternary uplift of the Kyrenia Range.

The structure of the foreland to the south of the Kyrenia Range can be inferred from several deep and shallow boreholes, combined with the available seismic data from beneath the Mesaoria (Mesorya) Plain. However, there is little relevant outcrop evidence except along the northern periphery of the Troodos Massif and adjacent to the Ovgos (Dar Dere) Fault.

A substantial thickness of Miocene sedimentary rocks (6–8 km) is inferred to exist between the Kyrenia Range front and the Ovgos (Dar Dere) Fault based on the prevailing stratal dip. Thrust repetition may be present at depth; alternatively, the area between the Ovgos Fault and the Kyrenia Range front could represent Miocene depocentres (effectively the site of a fossil trench).

The structure of the Mesaoria (Mesorya) Basin has previously been interpreted based on well and seismic data (Harrison et al. 2004a, b). Borehole data (Kato Lakatamia well) indicate the existence of stratigraphic repetition, which relates to southward-directed reverse faulting or thrusting (see Fig. 11). The borehole data also show that igneous rocks of the Troodos Massif extend northwards beneath the Mesaoria (Mesorya) Plain. Geophysical data also suggest that dense Troodos-like crust extends to near the front of the Kyrenia Range (Aubert and Baroz 1978). However, the deep crustal structure beneath the Mesaoria (Mesorya) Plain remains uncertain.

We are largely in agreement with the N–S cross section across the Mesaoria (Mesorya) Basin to the Kyrenia Range, as published by Harrison et al. (2004a, b). However, we integrate several modifications to this section in Fig. 11. First, field studies reveal the presence of important E–W, down-to-the-north high-angle faults cutting the sedimentary cover of the Troodos Massif, which affects the south of the section (Follows and Robertson 1990; Kinnaird and Robertson 2012). Secondly, biostratigraphic and strontium isotopic studies of pelagic carbonate rocks adjacent to the Ovgos (Dare Dere) Fault allow these to be assigned to a Miocene age (McCay and Robertson 2013) rather than to an older, Paleogene age. Thirdly, the large-scale structure of the Kyrenia Range has been modified taking into account of offshore seismic data (Calon et al. 2005a, b) and the results of this study.

Discussion and interpretation

Below, we integrate our new micro- to macro-scale structural and kinematic data with outcrop, to regional-scale tectono-stratigraphic information and use this to help infer the tectonic development of the central Kyrenia Range from Triassic to Pliocene. We also draw on the tectonic implications of sediment deposition and volcanism.

Deformation related to initial Late Cretaceous convergence

The Mesozoic shallow-water sedimentary rocks record accumulation on a subsiding carbonate platform (Robertson and Woodcock 1986). An extensional stress regime can be inferred for the Triassic, during a time of Neotethyan rifting. During the Late Cretaceous, the carbonate platform was deeply buried and underwent greenschist facies metamorphism in an inferred setting of regional convergence (Baroz 1979; Robertson and Woodcock 1986; Robertson et al. 2012b; Fig. 12a). The mostly bedding-parallel cleavage in the sparse mudrocks is likely to have originated during this event. However, other small-scale structures associated with early convergence appear not to have survived later deformation.

Fig. 12
figure 12

Cartoons summarising the inferred tectonic settings during Mesozoic-Recent time. a Late Cretaceous deep burial of the Mesozoic carbonate platform succession. This may have taken place in response to underthrusting of a carbonate platform unit, either a rifted continental fragment (Triassic) or a fragment detached from the southern margin of the Tauride continent; b exhumation and related brecciation of the Mesozoic platform carbonate; c Mid-Eocene southward thrusting (with an inferred left-lateral stress component), assumed to have taken place along an ocean-continent convergence zone; d Subsidence of the Kyrenia Range, possibly due to rollback of the S Neotethyan subduction zone; e Mid-Late Miocene renewed convergence, incorporation of the Kythrea (Değirmenlek) Group into the thrust stack; f latest Miocene-earliest Pliocene thrusting with a left-lateral stress component; g Pliocene, relative stability corresponding to a time of left-lateral strike-slip and ‘tectonic escape’ of Anatolia westwards towards the Aegean region; h Pleistocene, strong uplift influenced by collision of the Eratosthenes Seamount (a structural high) with the Cyprus trench to the south of Cyprus

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By Maastrichtian time, the platform carbonates were exhumed to the seafloor implying crustal extension (or transtension). The pervasive brecciation of the meta-platform carbonates could relate to this event. This was followed by erosion and mass wasting of the deformed meta-platform rocks in a subaqueous, marine, pelagic-depositing setting (Robertson et al. 2012b; Fig. 12b). It is tempting to relate the WNW–ESE- to ENE–WSW-trending sets of normal faults to the exhumation event, because such faults appear to be unrepresented within any of the younger units.

Latest Cretaceous-Eocene extension (or transtension)

During the latest Cretaceous (Maastrichtian) to Early Eocene, mass-flow units and calciturbidites were derived from underlying meta-platform rocks that were exposed as local subaqueous scarps (Robertson et al. 2012b). Texturally, this material is angular and immature, and there is no evidence that subaerial erosion was involved. Alkaline basaltic volcanic rocks of Maastrichtian and also Palaeocene-Early Eocene age erupted over the pelagic sediments, in places creating peperites. Their geochemistry indicates that the volcanism occurred in a within-plate, probably extensional (or transtensional) setting (Robertson and Woodcock 1986). The Campanian–Maastrichtian volcanism has also been related to back-arc rifting (Huang et al. 2007; Huang 2008) or the palaeo-rotation of the Troodos microplate (Robertson 1990). Regardless of the specific setting, the Kyrenia Range appears to have experienced a regional extensional (or transtensional) stress regime during Maastrichtian-Early Eocene time.

Mid-Eocene convergent phase

Thicker, more extensive mass-flow deposits appeared during the Early Eocene. These deposits include carbonate lithoclasts (Nummulitic limestone) and clasts of ophiolitic rocks (e.g. boninite) that are not known within the Kyrenia Range. Regional convergence resumed during the Early Eocene (Fig. 12c). Gravity-flow units (debris flows and turbidites) were deposited in a foredeep to the south of an advancing thrust load and also within subaqueous piggyback basins above the developing thrust belt. Up-to kilometre-sized blocks of Late Palaeozoic-Mesozoic shelf carbonate rocks were emplaced within sedimentary melange (~olistostrome) that is best developed in the eastern Kyrenia Range (Robertson et al. 2014).

The compression culminated in a major phase of ~east–west-orientated, mainly south-directed thrusting during the Mid-Eocene. A stack of gently inclined thrust sheets with associated folds was emplaced southwards, together with minor northeasterly to easterly verging folding. Intense shearing and pressure solution developed along the length of the central range during this convergent phase, as observed in the Palaeocene-Mid-Eocene pelagic carbonates (Ayios Nikolaos (Yamaçköy) Formation).

The observed component of westward displacement (Figs. 7, 9, 10) probably resulted from sinistral transpression. Limited additional kinematic data (e.g. from C–S fabrics) indicate a degree of westward displacement in agreement with sinistral transpression.

The NE–SW to ENE–WSW and the NW–SE to NNW–SSE faults occur in units that formed both before and after the Eocene convergence phase. However, the intensity and persistence of these fault trends are much greater within the Paleogene sedimentary rocks than within their Miocene and Pliocene counterparts. Such a contrast is clearly seen between the Palaeocene–Eocene Ayios Nikolaos (Yamaçköy) Formation and the Mid-Eocene Kalograia-Ardana (Baçeli-Ardahan) Formation (when taken together), compared with the unconformably overlying Late Eocene–Oligocene Bellapais (Beylerbey) Formation (e.g. near Beylerbey (Bellapais)). The structural differences between the rock units below and above this regional unconformity cannot be explained by rheological contrast alone (which is minor in some outcrops). Some of the NW–SE and the NNW–SSE-trending faults appear to have originated as transfer faults during the Mid-Eocene convergence phase and were later re-activated; numerous new faults of similar orientation were also created during the Late Miocene-earliest Pliocene convergence phase.

Stress regime during the Late Eocene-Late Miocene

The upward passage, over a short interval (<10 m), from coarse non-marine facies (conglomerates and sandstones) to finer-grained marine facies (sandstones and mudrocks) of the Late Eocene–Oligocene Bellapais (Beylerbey) Formation has been interpreted as the result of extension-related subsidence (Robertson and Woodcock 1986; McCay and Robertson 2012; Fig. 12d). Such extension could have been caused by southward rollback of a subducting slab during convergence along the northerly, active continental margin of the Southern Neotethys, of which the Kyrenia Range formed a part (Robertson et al. 2014).

To the south of the Kyrenia Range, the Kythrea (Değirmenlik) Fault was active during the Mid-Late Miocene, separating the Kyrenia Range and its southern periphery into northerly and southerly deep-water sub-basins in which contrasting Miocene siliciclastic gravity flows accumulated (Robertson and Woodcock 1986; McCay and Robertson 2012). The Kythrea (Değirmenlik) Fault is inferred to have existed as a northward-dipping sub-aqueous reverse fault until the Late Tortonian–Messinian when it was re-activated as a south-vergent thrust (McCay and Robertson 2012, 2013). Additional evidence from the southern part of Cyprus points to an overall extensional setting related to slab rollback within the Southern Neotethys during the Miocene (Kinnaird and Robertson 2012).

Late Miocene-earliest Pliocene convergence

Dominantly southward thrusting resumed during Late Miocene-earliest Pliocene time (Fig. 12e). The pre-existing south-directed Eocene thrust belt was re-activated and re-thrust further south, incorporating Late Eocene-Early Miocene siliciclastic sediments of the Bellapais (Beylerbey) Formation into the structurally lower levels of the thrust stack, as best developed in the south-western part of the central range (N of Selvili Tepe (Kornos)). In addition, the Ovgos (Dar Dere) Fault, further south, experienced compression and/or transpression during the Late Miocene-earliest Pliocene phase of convergence (McCay and Robertson 2013). The small-scale kinematic data from the Neogene sedimentary rocks of the southern flank of the Kyrenia Range (McCay and Robertson 2013) and the new data presented here are indicative of overall sinistral transpression during this time. The transpression also led to the formation, or reactivation, of transverse structures, including the eastern termination of the central range (near Antifonitis monastery), its western termination (near Selvili Tepe (Kornos); Fig. 4c, d), the formation of numerous large-scale oblique fault zones (e.g. Boğaz and Beşparmak (Pentadactylos) lineaments (Fig. 4e), numerous smaller-scale transverse faults [e.g. east of Lapta (Lapithos); Fig. 4b)] and ~N–S faults (Fig. 4f).

The major transverse lineaments do not appear to have existed, at least as significant bathymetric features, until the Late Miocene-earliest Pliocene phase of deformation. There is no evidence that the transverse features affected the deposition of the sediments of the Late Eocene-Late Miocene Kythrea (Diğirmenlik) Group. For example, very similar basal conglomerates are exposed within the Boğaz lineament (along the highway south of Kyrenia (Girne)) and both to the east (near Beylerbey (Bellapais)) and to the west (near Karaman (St Hilarion)). The major transverse structures can be traced northwards into the lower part of the Eocene-Miocene sedimentary succession, characterised by localised intense folding, shearing and soft-sediment deformation. The mapped NE-SW-trending left-lateral overlap of the Kyrenia Range stratigraphy (south of Kyrenia (Girne)) includes Late Eocene-Miocene sedimentary rocks, again showing that this feature post-dates Miocene sedimentation (Fig. 1).

The NNE–SSW, NW–SE and ENE–WSW-trending faults are intense along a few major oblique lineaments, but sparse along segments of the range in which the stratigraphy is relatively intact (e.g. Karaman (Karmi)). Some pre-existing ~N–S faults apparently were aligned in a favourable orientation to be reactivated and so accommodated some of the Late Miocene-earliest Pliocene deformation (Fig. 12f).

In summary, the Triassic-Early Eocene units of the central Kyrenia range were emplaced as a southward-verging pile of thrust sheets, although localised westward displacement took place, especially along the northern margin of the range, related to left-lateral oblique convergence. The steeply dipping to overturned nature of the unconformably overlying Late Eocene basal conglomerates implies that the thrust sheets were reactivated and steepened during the Late Miocene-earliest Pliocene phase of southward thrusting. During the latest Miocene-earliest Pliocene deformation phase, E–W sinistral transpression predominated, coupled with reactivation of NNW–SSE and NNE–SSW congruent faults. The deformation was accompanied by the development of a complex array of faults. Earlier generation faults were typically reactivated where in a favourable orientation to the later stress regimes.

Implications for regional tectonics

The structural development of the Kyrenia Range relates to stages in the evolution of the Southern Neotethys. A similar history is shown by the eastward extension of the Kyrenia Range into the Misis Mountain of coastal southern Turkey (Kelling et al. 1987; Gökçen et al. 1988; Robertson et al. 2004), although there are differences to the timing of tectonic events in detail.

Following rifting during the Triassic and subsequent Jurassic-Cretaceous passive margin subsidence (Robertson and Woodcock 1986), regional convergence was initiated during the Late Cretaceous. The Kyrenia platform represents the southernmost part of the Tauride continent or a more southerly rifted continental fragment within the Southern Neotethys (Robertson et al. 2012b, 2014). The Kyrenia platform was thrust beneath its Tauride continental backstop to the north, apparently as an intact structural unit, and then exhumed by Maastrichtian time. Northward subduction gave rise to calc-alkaline arc-type volcanism, including rhyolitic tuffs and minor intrusions, during the Campanian–Maastrichtian (?), as best exposed in the western Kyrenia range (Baroz 1980; Huang et al. 2007; Robertson et al. 2012b; Gilbert and Robertson 2013). Maastrichtian-Early Eocene time was characterised by two phases of within-plate basaltic volcanism in an inferred extensional (or transtensional) setting.

Southward convergence during the Mid-Eocene resulted in the large-scale emplacement of mass-flow units of the Kalograia-Ardana (Baçeli-Ardahan) Formation (Robertson et al. 2014). The convergence culminated in southward thrusting of the Triassic-Eocene units making up the Kyrenia Range. The available kinematic data (e.g. C–S fabrics and asymmetric fold trends) indicate overall compression/sinistral transpression during Mid-Eocene time. Some of the ~N–S transverse, high-angle faults are likely to have developed as linking transfer structures during this stage. The likely cause of the southward thrusting was a resumption or acceleration of northward subduction of the Southern Neotethys, accompanying or following the final closure of the Izmir–Ankara–Erzincan ocean further north (see Robertson et al. 2014).

Late Eocene-Late Miocene time was characterised by the southward retreat of the South Neotethyan subduction zone, resulting in subsidence. Related to on-going convergence, the incoming plate segmented into two ~E–W depocentres separated by the Kythrea (Değirmenlik) Fault. Southward thrusting was again active during the Late Miocene-earliest Pliocene, in response to suture zone tightening, as documented in SE Turkey (Robertson et al. 2004). As a result, the pre-existing Kyrenia Range thrust belt was re-thrust southwards within an overall sinistral transpressive regime. The sinistral component of thrusting triggered the development of numerous ~N–S transfer faults. The associated oblique faults are interpreted as the result of the overall sinistral transpressive stress regime.

The Pliocene marine sedimentary rocks accumulated in a relatively quiescent tectonic setting during the westward tectonic escape of Anatolia to the north, bounded by sinistral faults, including the Ovgos (Dar Dere) Fault in Cyprus (McCay and Robertson 2012; Fig. 12g).

Strong uplift of the convergent margin ensued during the Late Pliocene-Early Pleistocene possibly influenced by the collision of the Eratosthenes Seamount with the Cyprus trench to the south of Cyprus (Robertson 1998b; Kinnaird et al. 2011; Fig. 12h). The varied Quaternary shallow-marine to continental deposits flanking the Kyrenia Range show little evidence of faulting or folding (Palamakumbura et al. 2015).

The tectonic development of the central Kyrenia Range shows similarities and also differences compared to the equivalent development of the Southern Neotethys in SE Turkey. Late Cretaceous thrusting and subduction-related magmatism are important features of both regions (e.g. Nurlu et al. 2015). Convergence along the northern margin of the Southern Neotethys appears to have been on-going at least during Eocene–Oligocene in SE Turkey (e.g. in the Engizek Mountains; see Robertson et al. (2004) and Akınçı et al., this volume). The Southern Neotethys sutured in SE Turkey during the Early Miocene, during a time when deep-water sedimentation persisted in the Kyrenia Range. However, the convergence-related Kythrea (Değirmenlik Fault) may relate to the regional collision-related crustal deformation. The suture zone in SE Turkey experienced southward re-thrusting coupled with folding (suture zone tightening) during the Late Miocene-Early Pliocene, during the same time as re-thrusting and folding in the Kyrenia Range.

The differences in the tectonic development of the two adjacent regions (N Cyprus and SE Turkey) relate to the diachronous closure of the Southern Neotethys. Relatively early and much more pervasive continental collision took place in the SE Turkey compared to the Kyrenia Range. The likely explanation is that the Arabian promontory collided with the Tauride continent during the Early Miocene in SE Turkey, whereas the Kyrenia Range (and Cyprus as a whole) remained within a relict oceanic basin to the west. The main uplift of both the Troodos Massif in south Cyprus and the Kyrenia Range is likely to have been controlled by the collision of the Eratosthenes Seamount with the Cyprus active margin in the south, still within the relict ocean basin.

Conclusions

The structural development of the central Kyrenia Range reflects stages in the development of the Southern Neotethys, specifically Late Cretaceous to Late Miocene subduction and diachronous collision. Initial convergent deformation during the Late Cretaceous resulted in greenschist facies metamorphism, followed by exhumation. A second period of convergent deformation took place during the Middle Eocene giving rise to southward thrusting that was coupled with sinistral transpression and the formation of ~N–S transverse (transfer) faults, oblique faults and conjugate faults. A third period of convergent deformation occurred during Late Miocene (Messinian)-earliest Pliocene time giving rise to southward thrusting and localised back-thrusting, coupled with variable-scale folding and common fault reactivation and also large-scale segmentation of the range associated with ~N–S transverse (transfer) faulting and oblique ~NE–SW/NW–SE faulting.

Regional comparisons suggest that northward subduction along the northern margin of the Southern Neotethys was active during Late Cretaceous to Miocene time (although probably with changing kinematics). The Late Cretaceous initial convergent event relates to under-thrusting of a carbonate platform unit beneath the Tauride continent. The Mid-Eocene convergent event relates to the focussing of northward subduction within the Southern Neotethys after other ocean basins in central Turkey closed. To the east, the Arabian peninsula collided with the Tauride continent during the early Miocene, whereas Cyprus remained in the by-then relict Southern Neotethys that evolved into the Eastern Mediterranean basin. However, the regional-scale collision influenced Mid-Late Miocene sedimentation in and around the Kyrenia Range. The latest Miocene-earliest Pliocene convergent event in the Kyrenia Range corresponds to a phase of suture zone tightening (thrusting and folding) in SE Turkey, prior to onset of the westward ‘tectonic escape’ of Cyprus towards the Aegean region. Strong uplift of the Kyrenia Range as a whole took place during the Late Pliocene–Pleistocene, possibly controlled by the collision of the Eratosthenes Seamount (a submerged platform) with a subduction trench to the south of Cyprus.