The Kyrenia Range underwent tectonically driven uplift during the Pliocene to Pleistocene in response to the interaction of various tectonic processes. To understand the tectonic processes driving the uplift and how this is related to uplift of other areas of the Eastern Mediterranean, uranium-series disequilibrium and optically stimulated luminescence dating were applied to marine and non-marine terrace deposits exposed on the northern flank of the range. Palaeomagnetism and strontium isotope dating were used in conjunction to date the final stages of the marine environment adjacent to the Kyrenia Range prior to major surface uplift. Uplift rates range from >1.2 mm a−1, inferred during the Early Pleistocene, to <0.2 mm a−1 during the Late Pleistocene. The new data show that the Kyrenia Range was uplifted contemporaneously with the Troodos Massif in southern Cyprus. The uplift of the Kyrenia Range appears to have been significantly faster than that affecting other comparable regions in the easternmost Mediterranean during the Pleistocene (e.g. Lebanon coast; southern Anatolian plateau). The driving mechanism for the uplift of both the Kyrenia Range and the Troodos Massif is inferred to be the collision of the Eratosthenes Seamount with the Cyprus trench to the south of Cyprus.
Supplementary material: Details of the methods used are available at https://doi.org/10.6084/m9.figshare.c.3260977.
The main aim of this paper is to present and discuss the results of quantitative dating of Pliocene and Pleistocene deposits associated with the Kyrenia Range with a view to understanding the timing, rates and processes driving its surface uplift. Similar studies have been carried out on other areas, such as the Miocene to Recent uplift of the Andes (Gregory-Wodzicki 2000) and the Late Cenozoic uplift of southeastern Tibet (Clark et al. 2005). A key current objective of geological and geomorphological research is the identification of the crustal processes involved in surface uplift. The Eastern Mediterranean region is very well suited to this approach because it is an active tectonic area that is becoming increasingly well understood based on a combination of both land and marine geophysical and geological studies. The easternmost Mediterranean region (Fig. 1) is a convergence zone involving subduction, strike-slip (including transpression and transtension) and collision at different scales, together with deep-seated processes such as crustal delamination, slab break-off and slab tearing (Faccenna et al. 2006; Berk Biryol et al. 2011; Schildgen et al. 2012; Bartol & Govers 2014). The interaction of such deep-seated tectonic processes drives surface uplift in various parts of the easternmost Mediterranean and the adjacent region. For example, major surface uplift occurred during the Plio-Pleistocene in the Peloponnese and elsewhere in Greece (Roberts et al. 2009; Turner et al. 2010) and in the Taurus Mountains of southern Turkey (Schildgen et al. 2012).
Previous work has documented the uplift of the Troodos Massif in the south of Cyprus, which climaxed during the Pleistocene (Gass 1960; De Vaumas 1961; Poole et al. 1990; Poole & Robertson 1991, 2000; Robertson & Xenophontos 1993; Kinnaird et al. 2011; Kinnaird & Robertson 2013; Main et al. 2016). The Troodos Massif is a Late Cretaceous ophiolite (subaerially exposed oceanic crust), overlain by Late Cretaceous to Recent deep-marine to shallow-marine and non-marine sediments. The likely cause of the uplift of the Troodos Massif was the collision of the Eratosthenes Seamount, a large crustal edifice, with the Cyprus active margin, related to a northward-dipping subduction zone. The uplift was possibly localized and accentuated by the effects of serpentinization of the oceanic mantle wedge by fluids derived from the downgoing slab (Robertson 1990, 1998b, 2000; Kempler 1998).
Sandwiched between southern Turkey and the Troodos Massif is the Kyrenia Range, an approximately east–west-trending narrow, arcuate mountain range in the northern part of Cyprus. The western end of the Kyrenia Range reaches c. 1000 m above modern sea level (AMSL); the range decreases in height eastwards overall and then passes into a submerged ridge that links with the Misis Mountains of southern Turkey (Fig. 1). Previous studies have suggested that the Kyrenia Range is of relatively recent origin (Ducloz 1972; Baroz 1979) but the detailed timing and processes involved remained elusive. New evidence for the uplift is presented here in relation to adjacent areas and alternative interpretations are discussed in their regional context.
The regional geology of the area concerned may influence the nature of any uplift, including features such as the relative competency of lithologies and inherited structure.
In outline, the Kyrenia Range has a Late Palaeozoic to Recent geological history with four main stages of stratigraphical development. These encompass the deposits of the Trypa (Tripa), Lapithos (Lapta), Kythrea (Değirmenlik) and Mesaoria (Mesarya) Groups (it should be noted that the nomenclature follows stratigraphic priority; Fig. 2). We use the traditional stratigraphic names, followed by recent Turkish equivalents. The Trypa (Tripa) Group is dominated by neritic metacarbonates of Triassic to Cretaceous age (Ducloz 1972; Baroz 1979; Robertson & Woodcock 1986). The Lapithos (Lapta) Group unconformably overlies the Trypa (Tripa) Group and comprises pelagic carbonates interbedded with subaqueous basic or silicic volcanic rocks and also sedimentary mélange (olistostromes), of overall Maastrichtian to Eocene age (Robertson et al. 2012). The unconformably overlying Kythrea (Değirmenlik) Group comprises Late Eocene to Late Miocene, mostly marine conglomerates, sandstones, marls and mudstones, culminating in Messinian evaporites (Weiler 1970; McCay 2010; McCay & Robertson 2012a; McCay et al. 2012; Taylforth et al. 2014). Unconformably above comes the Mesaoria (Mesarya) Group, represented by Pliocene shallow-marine to non-marine sediments (Ducloz 1972; Baroz 1979; McCay & Robertson 2012a; McCay et al. 2012). The Mesaoria (Mesarya) Group includes the Myrtou (Çamlıbel), Nicosia (Lefkoşa), Athalassa (Gürpınar) and Fanglomerate Formations (Baroz 1979). Of these, the Myrtou (Çamlıbel) Formation comprises interbedded marls and siltstones (Baroz 1979; Yetis et al. 1995; McCay et al. 2012), whereas the Nicosia (Lefkoşa) Formation is made up of interbedded marls and chalks (Baroz 1979; Yetis et al. 1995). Above this, the Athalassa (Gürpınar) Formation is predominantly shallow-marine carbonate grainstones (Baroz 1979; McCallum & Robertson 1995a; Harrison et al. 2004). Finally, the traditional Fanglomerate Formation is represented by a series of terrace deposits on both flanks of the Kyrenia Range (Ducloz 1964; Knup 1965; Baroz 1979).
The opposing flanks of the Kyrenia Range are characterized by terraced littoral-marine and non-marine sediments, up to a height of c. 180 m AMSL. These deposits were previously mapped and described as being Pleistocene (Ducloz 1972; Dreghorn 1978), although no absolute age data were available. Based on extensive new fieldwork (Palamakumbura 2015; Palamakumbura et al. 2016), five main terraces are now identified on the northern and the southern flanks of the range, defined as K0–K5 (Figs 2 and 3a). The K0 terrace is the highest and, by inference, the oldest terrace, whereas the K5 terrace is the lowest and youngest terrace. In terms of the previous stratigraphy the K0, K1, K2, K3, K4 and K5 terraces are equivalent to the Karka, Klepini (Arapköy), Trapeza (Beşparmak), Ayios Epikititos (Çatalköy), Kyrenia (Girne) and Koupia terraces, respectively (Ducloz 1972). The highest terrace, K0, is preserved at c. 800–600 m AMSL. The K0 terrace comprises non-marine megabreccia with no preserved marine facies. The K1 terrace is preserved on the distal parts of the K0 terrace at 600 m AMSL and comprises colluvial conglomerates. The K2, K3 and K4 terrace deposits, which include shallow-marine facies, are preserved at c. 180–140, 70–50 and 40–0 m AMSL, respectively. Finally, the K5 terrace is preserved along the northern coast at modern sea level. The age progression of the terraces has been inferred from outcrop field relations, including incision and facies geometry, and was also tested by the age dating presented here. The heights of each terrace are taken from the most northerly part of each terrace, closest to the coast, reflecting the basal and oldest part of each terrace system. The height above sea level of each terrace was established from a combination of the contoured base of the published geological maps of the central Kyrenia Range (Ducloz 1964; Knup 1965), handheld global positioning system (GPS) measurements and Google Earth. The estimated error in the heights is <5 m, whereas the variation in height between each of the terrace systems (other than the two youngest ones) is several tens of metres or more. Uncertainties in absolute height do not compromise our discussion of the processes of uplift v. sea-level change.
The structure of the Kyrenia Range and the adjacent Mesaoria (Mesarya) Basin has been documented based on regional mapping (Baroz 1979) from outcrop-scale observations and microstructural studies (McCay & Robertson 2012a; Robertson & Kinnaird 2016). The last of several major phases of deformation corresponds to an important phase of southward thrusting during Late Miocene–earliest Pliocene time in a left-lateral, transpressional stress regime (the earlier structural history is not considered here). Pleistocene compression or transpression is locally documented by faults that cut Pliocene sediments (McCay & Robertson 2012b). In addition, a small number of transverse fault planes within the Kyrenia Range are mantled by relatively immature scree talus, of inferred Pleistocene age (e.g. NNW–SSW-trending fault lineament near Beşparmak (Pentadactylos mountain) (Robertson & Kinnaird 2016)). There is also evidence of Pliocene–Pleistocene sinistral strike-slip movements along the Ovgos (Dar Dere) Fault to the south of the Kyrenia Range (Harrison et al. 2004).
Here, we combine field-based studies of the Pliocene and Pleistocene sediments with absolute age dating using several well-established techniques, each of which have pros and cons. Classical biostratigraphy (e.g. using macrofossils) is of limited use to date the Pliocene–Pleistocene sediments because most of the species present are relatively long ranging. All but possibly the very youngest sediments are expected to be beyond the age range of radiocarbon dating. Optically stimulated luminescence (OSL) is potentially useful, but only for sediments of <200 ka age (for quartz) and <500 ka age (for feldspar) (Rhodes 2011). Similarly, uranium-series disequilibria (U-series) dating is unlikely to be viable at >350 ka (Van Calsteren & Thomas 2006). U-series dating depends on analysing aragonitic coral, which although present is relatively rare. Other macrofossils including bivalves rarely provide reliable results (Van Calsteren & Thomas 2006). Magnetostratigraphy has obvious potential, following successful studies of similar deposits in the south of Cyprus (Kinnaird et al. 2011; Weber et al. 2011), but is dependent on the sediments preserving measurable stable magnetic remanence that is indicative of magnetic polarity (Kodama 2012). Strontium isotopic analysis is also useful but only for the dating of suitable, well-preserved fossils from an open-marine setting (Beets & De Ruig 1992; Müller 1993). Exposure dating could also have potential, although detailed geomorphological analysis was outside the scope of this work.
In practice, some valid new data were obtained with each of the techniques used, although all of these are problematic in one way or another. The techniques used followed well-established methods. U-series and OSL dating proved suitable for dating of the younger marine terraces (K4 and K5) on the northern flank of the range. In addition, microfossil evidence, mostly from planktonic foraminifera, and palaeomagnetic data were effective for dating shallow-marine sediments that were deposited before major uplift of the range.
Uplift deposits and dating
U-series dating was carried out on well-preserved solitary corals (Cladocora caespitosa) (Fig. 3b and c). The corals were collected from the K4 terrace along the northern coast. The best-preserved coral samples came from two marine deposits, as summarized in Table 1. The first dated deposit is a marine conglomerate, which includes well-rounded and well-sorted clasts of locally derived ‘basement’ rocks (e.g. metacarbonate rocks, basalt, chert), ranging in size from 1 to 10 cm. Large amounts of biogenic material also occur within the conglomerate including oysters, pectens, gastropods, serpulid worms (as casts), molluscs (commonly bored) and solitary coral. The second dated deposit is a breccia of well-cemented marine bioclastic grainstone (Table 1). This is fine to medium grained and contains abundant reworked biogenic material including molluscs, serpulid worm casts, bored clasts and solitary coral. The observed marine grainstone is interbedded with a Gilbert-type deltaic deposit.
OSL dating was carried out on aeolian grainstones of the K5 terrace on the northern coast (Fig. 3d and e). The grainstones (Table 1) are medium grained and contain well-sorted and well-rounded grains. The coastal grainstones contain well-preserved trough cross-bedding, with 1–3 m thick foresets. Calcified palaeo-root traces are well preserved throughout the aeolian deposits. In a related study, the grainstone deposits have been profiled in the field using a portable OSL reader to assess their luminescence characteristics (Palamakumbura et al. 2016). This study showed that these sediments contain sufficient quartz and feldspar with sufficient luminescence signal to permit dating. The deposit is interpreted as representing a coastal aeolian dune environment (Palamakumbura 2015; Palamakumbura & Robertson 2016; Palamakumbura et al. 2016).
Uranium-series disequilibrium dating
Solitary corals are relatively rare in the littoral marine terrace deposits along the northern coast. Also, some of those collected were found, based on X-ray diffraction (XRD) analysis, to be partially recrystallized and so unsuitable for dating (Palamakumbura 2015). Following extensive fieldwork, five solitary coral samples were selected for dating from four sites within the K4 terrace along the northern coast (Fig. 2). The selected corals are preserved in situ without reworking and are, therefore, believed to represent the age of the marine deposit in which they are preserved. The coral samples were cut into several subsamples, with the central septa being removed and the walls cleaned with a Dremel drill©. XRD analysis was carried out on the coral-wall samples. Only coral samples with over 97% aragonite were used for dating.
The facies of the deposits containing the corals is indicative of a major marine transgression event followed by a marine regression. In addition, relatively minor transgressive events are indicated by the presence of thin (tens of centimetres) layers of coral-bearing deposits within otherwise regressive sedimentary sequences, as discussed in detail elsewhere (Palamakumbura 2015; Palamakumbura & Robertson 2016; Palamakumbura et al. 2016).
The U-series dating was carried out at the Open University, Milton Keynes, using standard methods: further information has been given by Van Calsteren & Thomas (2012). Samples KY12-03, KY12-04 and KY11-22 have 230Th/234U ratios of <1 (Table 2). In contrast, samples KY13-19(b), -94(a) and -94(b) exhibit 230Th/234U ratios of c. 1, indicating either that the system has reached secular equilibrium or that thorium isotopic contamination has taken place. The 232Th concentrations are low enough to suggest that thorium contamination is not a significant issue for any of the samples (Table 2). The samples KY13-19(b), -94(a) and -94(b) are, therefore, assumed to be in secular equilibrium. Table 2 shows the final calculated ages of the coral samples, as determined from the relationship between uranium and thorium isotopes (Ivanovich 1982). The five coral samples yielded ages ranging from 127 to >350 ka (Middle Pleistocene).
Optically stimulated luminescence
The luminescence dating analysis was carried out at the Scottish Universities Environmental Research Centre (SUERC) luminescence laboratories, East Kilbride. Conventional OSL quartz and also post-infra-red elevated temperature infra-red-stimulated luminescence (i.e. post-IR IRSL) potassium feldspar approaches were initially considered as potential methods for dating the aeolianite. The regional geology is dominated by calcareous rocks of Permian to Middle Miocene age, together with subordinate volcanogenic and siliceous rocks. This is challenging because locally derived sands are largely composed of low-radioactivity materials with little quartz or alkali feldspar. However, luminescence profiling, both during the fieldwork and subsequently in the laboratory revealed stratigraphically progressive OSL and IRSL signals (Palamakumbura et al. 2016), indicating the presence of material in the sediments with dating potential. Although the laboratory characterization indicated the presence of some quartz, its low yields and low luminescence sensitivity, and (on further investigation) the lack of thermal stability of the conventional OSL signals (i.e. the fast component) precluded application of quartz-based methods. Electron microscopy confirmed the presence of feldspars, which were separated and used instead of quartz for IRSL and elevated temperature post-IR IRSL single aliquot regeneration (SAR) dating (Buylaert et al. 2009; Thomsen et al. 2011; Kinnaird et al. 2015).
Dose-rate estimates from the carbonate sands were assessed using a combination of field gamma spectrometry, high-resolution gamma spectrometry (HRGS) and thick-source beta counting (TSBC), which were reconciled with each other and with the water contents and micro-dosimetry of the model. In addition to the two dating samples, a series of eight bulk gamma spectrometry samples were collected to permit retrospective analysis of the external gamma dose rates. The gamma dose rates were reconstructed from seven of these, giving wet gamma dose rates of between 0.4 and 0.7 mGy a−1, which are comparable with the measured values recorded at the sampling positions (0.4–0.5 mGy a−1). Counting times for both HRGS and TSBC measurements were extended by one order of magnitude, resulting in overall uncertainties of <4% for these low dose rates. Internal alpha dose rates were estimated on the basis of inductively coupled plasma mass spectrometry (ICP-MS) analysis of comparable feldspars and make a minor contribution to the overall dose rates from this material.
Equivalent doses were determined from 16 aliquots of K-feldspar per sample using IRSL and elevated temperature post-IR IRSL SAR protocols. The material exhibited good IRSL sensitivity and produced acceptable SAR internal quality control performance. On the basis of these results, coupled with the IRSL SAR quality selection criteria and the expectations of the procedure, preference was given to the post-IR IRSL 225°C signals in equivalent dose determinations (Fig. 4). Radial plotting methods revealed some heterogeneity in the equivalent dose distributions of each sample, reflecting variable bleaching at the time of deposition and indicating that each sample enclosed mixed-age materials (Fig. 4c). The laboratory profiles provide some measure of control on this. The basal parts of each deposit show the greatest dispersion or spread in IRSL stored dose values, and, importantly, the lowest values in each profile. In contrast, the upper units are characterized by larger IRSL stored dose values with tighter distributions, reflecting faster deposition and poor to incomplete bleaching, or zeroing, of the luminescence signals. The mean ratios of sensitization-independent stored to prompt normalized signals after 106 s storage are 0.93 ± 0.06 and 1.08 ± 0.08, for IRSL and post-IR IRSL 225°C, respectively.
Luminescence ages were calculated using standard microdosimetric models, with uncertainties that combined measurement and fitting errors from the SAR analysis, all dose rate evaluation uncertainties, and allowance for the calibration uncertainties of the sources and reference materials. The post-IR IRSL 225°C age estimates are (SUTL2698) 76.1 ± 3.5 ka and (SUTL2699) 58.7 ± 2.6 ka (Table 3).
The OSL dates from the K5 terrace come from aeolian grainstones along the northern coast. Both of the dated deposits are preserved at modern sea level but are 30 km apart along the coast. These deposits are interpreted as two distinct aggradation events at 76.1 and 58.7 ka during marine isotope stage (MIS) 4. Both deposits are interpreted as belonging to the K5 terrace and are likely to record different depositional phases within a single terrace system.
Age of pre-uplift deposits
An additional aim of this study was to achieve a better understanding of the timing of deposition prior to the first major uplift of the Kyrenia Range. Dating of the pre-uplift deposits was previously achieved through a combination of foraminiferal biozonation (Baroz 1979; Hakyemez et al. 2000; A. Lord, cited in Harrison et al. 2008) and limited strontium isotopic dating (McCay et al. 2012).
Deposition prior to major uplift is represented by the Nicosia (Lefkoşa) and Athalassa (Gürpınar) Formations (Ducloz 1972; Baroz 1979). The Nicosia (Lefkoşa) Formation comprises interbedded chalks and marls, whereas the Athalassa (Gürpınar) Formation is dominated by neritic grainstone facies (Palamakumbura & Robertson 2016) (Table 1; Fig. 3f). Here we use the traditional classification of the Athalassa (Gürpınar) Formation by Ducloz (1972). Dating of these deposits has been achieved using a combination of foraminiferal biozonation, palaeomagnetism and strontium isotope dating. The Nicosia (Lefkoşa) and Athalassa (Gürpınar) Formations are exposed within the Mesaoria (Mesorya) Basin between the Kyrenia Range and the Troodos Massif (Fig. 2). During this work samples were collected from the northern part of the Mesaoria Basin, close to the Kyrenia Range (north of the Ovgos (Dar Dere) Fault).
Strontium (87Sr/86Sr) dating
Strontium isotope dating is based on comparing the 87Sr/86Sr ratio from a sample with 87Sr/86Sr ratios in global oceans through time, expressed as a reference curve (McArthur et al. 2001). Two types of sample were used for the present strontium analysis, oyster shells and bulk grainstone (Table 4). Oysters were used as they are abundant and relatively well preserved, in contrast to most other bivalves in the deposits. The samples were collected from the upper part of the Athalassa (Gürpınar) Formation (Table 4). The oyster shells appear to be well preserved and in life position within benthic foraminifera-rich grainstones. Preparation of the oyster shells for analysis was carried out using the method of Henderson et al. (1994), whereas the preparation of the bulk grainstone followed that of Bailey et al. (2000). The strontium analysis was carried out at SUERC, East Kilbride on a VG Sector 54-30 mass spectrometer in a dynamic multi-collection mode.
Based on point counting, the grainstones analysed for strontium dating comprise abundant benthic foraminifera (c. 54%), mostly Peneroplidae, together with planktic foraminifera (c. 12%), reworked echinoderms (c. 4%), calcareous red algae (c. 5%), bivalve shell fragments (c. 2%) and reworked siliciclastic material (c. 21%). The deposits include finely crystalline calcite spar cement that coats the grains and partially infills the pore space.
The oyster shells and the bulk grainstone yielded 87Sr/86Sr ratios ranging from 0.70872 to 0.70911, representing an inferred age range of 15.99–1.43 Ma (Middle Miocene to Middle Pleistocene).
The Miocene ages results are problematic because the Athalassa (Gürpınar) Formation is known to unconformably overlie the Late Eocene–Late Miocene Kythrea (Değirmenlik) Group, which in places terminates with gypsum that was precipitated during the Messinian salinity crisis (Baroz 1979; McCay et al. 2012). It seems, therefore, that the Miocene Sr isotopic ages have to be discounted. Possible reasons include diagenetic alteration of the oyster shells, reworking of older material or deposition in a semi-enclosed marine basin with abnormal strontium isotopic ratios. Reworking is unlikely because the oyster shells appear to be in life position within shallow-marine deposits. The samples dated as Miocene are from the central part of the Mesaoria (Mesarya) Basin, which was a narrow seaway between the Troodos and Kyrenia landmasses during Late Pliocene–Early Pleistocene time (McCallum & Robertson 1995b) in which salinity could have been anomalous. The most likely explanation for the age discrepancy is that the oyster shells were affected by diagenesis that altered the 87Sr/86Sr ratios. The one whole-rock sample that was dated (sample KY13-80) came from the Karpaz (Karpas) Peninsula in the far east of the Kyrenia Range (Fig. 2), an area with a more open-marine influence. The determined Early Pleistocene age is consistent with the expected age of the Athalassa (Gürpınar) Formation, compared with the known Pliocene age of the underlying Nicosia (Lefkoşa Formation), as discussed below. However, no reliance can be placed on a single age result.
Palaeomagnetic dating has previously been used to help constrain the age of Pliocene and Pleistocene shallow-marine sediments around the southern and eastern margins of the Troodos Massif (Kinnaird et al. 2011; Weber et al. 2011). A similar approach was adopted for the sedimentary rocks exposed within and adjacent to the Kyrenia Range during this study. Samples were collected using a portable petrol drill and oriented using a sun compass, according to standard field methods (Butler 1992). After conventional measurement and analysis, it had to be concluded that none of the samples from the relatively young (‘Pleistocene’) terrace deposits on either flank of the Kyrenia Range yielded viable palaeomagnetic results (Palamakumbura 2015). However, the pre-uplift sequence exposed on the outer southern flank of the range provided worthwhile palaeomagnetic data.
The successful palaeomagnetic analysis was carried out on shallow-marine grainstones of the Athalassa (Gürpınar) Formation. Ninety-five cores from 19 sites were collected from the Nicosia (Lefkoşa) Formation and the overlying Athalassa (Gürpınar) Formation. All of the samples were first studied utilizing alternating-field (AF) demagnetization techniques at the National Oceanography Centre, University of Southampton and were measured using a 2G Enterprise Superconducting Rock Magnetometer. This revealed that 50 of the cores had a sufficiently intense magnetic signal to warrant thermal demagnetization, which was carried out on sister samples using a 2G Enterprises DC Squid Cryogenic Magnetometer at the University of Oxford. The thermal demagnetization steps used 50°C steps from 0 to 300°C, followed by 30°C steps, up to 590°C. Assessment of the quality of the data was mainly based upon inspection of the resulting Zeijderveld plots and from the intensity-decay curves (Zijderveld 1967). Each sample was ranked from A to C using the system developed by Richter et al. (1998), with A representing samples with a clear stable magnetic direction and C samples with either unstable directions or no magnetic signal. The A and B quality data were selected for principal component analysis (PCA).
When analysed using conventional methods (Butler 1992) the stepwise thermal demagnetization experiments revealed stable components of magnetization in 15 of the grainstone samples (Fig. 5 and Table 5). These samples were selected for magnetostratigraphic analysis. Figure 5 shows orthogonal projections of the thermal demagnetization data. The samples show a general reduction in magnetic signal intensity with increasing temperature. Removal of an unstable low-temperature signal (20–400°C) revealed a stable high-temperature direction of demagnetization between 400 and 590°C. This high-temperature style end-point direction was used to calculate the primary palaeomagnetic direction for each sample (Table 5).
Although the primary magnetization could be isolated in only 15 samples, it is clearly identified as having a dual polarity within the Athalassa (Gürpınar) Formation, with samples from the lower part of the sequence (sample PM12-08 A, B, D, E and PM12-10 A, B, C, D; Fig. 5 and Table 5) being of normal polarity (northward with positive inclinations) whereas the samples from the higher part of the sequence (PM12-09 A, B, C, D, E; Fig. 5 and Table 5) are of inverse polarity (southward and negative). The low α95 and high κ values for sites PM12-09 and PM12-10 (Table 5) are indicative of reasonably robust data. Site PM12-18 has relatively high α95 and low κ values indicating a less reliable dataset.
The large scatter of inclination values is possibly due to local sedimentary processes such as current orientation of grains (Kodama 2012). Sediment compaction was previously inferred to have affected measured inclination values in the Pliocene–Pleistocene sediments from the southern part of Cyprus (Kinnaird et al. 2011).
Despite the limitations of data quality, the palaeomagnetic results from the Athalassa (Gürpınar) Formation suggest that the deposition spanned a geomagnetic reversal. Reversed polarity grainstone deposits overlie normal polarity grainstones. The reversed polarity grainstones are interpreted as having been deposited before the Brunhes Chron, which marks the beginning of the Middle Pleistocene (Fig. 6). We have no constraints on the specific chrons with which these deposits could correlate, although an Early Pleistocene age is likely in view of the Pliocene age of the underlying Nicosia (Lefkoşa) Formation (see below).
Timing of uplift
The viable results from the different dating techniques are shown in Figure 6 in relation to the Pliocene–Pleistocene stratigraphy of the Kyrenia Range. Mapping and sedimentological studies around the southern flank of the Kyrenia Range indicate that the entire intact Pliocene succession predates the oldest marine terrace deposits (K2), which occur at heights of up to 160 m AMSL (Palamakumbura 2015; Palamakumbura & Robertson 2016).
The chalks and marls of the Nicosia (Lefkoşa) Formation are rich in planktonic foraminifera, including Globorotalia magaritae and Globorotalia acostaensis (Baroz 1979; Hakyemez et al. 2000; A. Lord, pers. comm. 2015). The biota identified within the Nicosia (Lefkoşa) Formation can be interpreted as Early to Late Pliocene age based on up-to-date statigraphical correlations with oceanic deposits (Lourens et al. 2004; Wade et al. 2011). Several boreholes in the Greater Lefkoşa (Nicosia) area provide additional data on the age of the Nicosia (Lefkoşa) Formation. The planktonic foraminiferal assemblages are indicative of zones N19 to N20, of Pliocene age (A. Lord, cited by Harrison et al. 2008). The pre-uplift open-marine environment represented by the Nicosia (Lefkoşa) Formation is, therefore, inferred to have existed during the Pliocene. It should be noted that sediments of the Mesaoria (Mesorya) Group are not preserved along the northern flank of the Kyrenia Range although they may originally have been deposited there.
The Athalassa (Gürpınar) Formation conformably overlies the Nicosia (Lefkoşa) Formation within the northern part of the Mesaoria (Mesarya) Basin (Baroz 1979; this study). Similar planktonic foraminifera to those within the Nicosia (Lefkoşa) Formation have previously been reported from the Athalassa (Gürpınar) Formation in this area (Baroz 1979; Hakyemez et al. 2000). The Athalassa (Gürpınar) Formation is dominated by carbonate grainstones that accumulated in a shallow-marine (sub-littoral) environment (Baroz 1979; McCallum & Robertson 1995b; Harrison et al. 2004). Petrographic study of the grainstone deposits indicates that the formation includes large amounts of reworked shallow-water bioclastic material and also reworked planktonic foraminifera (Palamakumbura 2015). Reworking of planktonic foraminifera is suggested by (1) broken or abraded tests, (2) infilling of chambers with a different diagenetic cement from that of the matrix and (3) local presence of pre-Pliocene species. Some of the planktonic foraminifera that have previously been identified from the Athalassa (Gürpınar) Formation (Baroz 1979; Hakyemez et al. 2000) are, therefore, likely to have been reworked from the Nicosia (Lefkoşa) Formation. However, some planktonic foraminifera appear to be contemporaneous; these are preserved as intact tests, which enclose fine-grained micrite of similar composition to the sedimentary matrix. The shallow-marine deposits of the Athalassa (Gürpınar) Formation accumulated during a period of, first reversed polarity, then normal polarity, which is consistent with an Early Pleistocene age.
The U-series dating constrains the age of several littoral marine horizons related to the K4 terrace. When plotted on a eustatic sea-level curve, utilizing a widely used oxygen isotope stratigraphy (Siddall et al. 2006), the absolute age results correspond to two interglacial periods, ranging from MIS 7 to MIS 5 (Fig. 6). The >350 ka U-series date is believed to be analytically valid but is difficult to interpret, because the coral in this terrace deposit is fragmentary and reworked.
In addition, the OSL data constrain the age of the aeolian deposits associated with the K5 terrace (Fig. 6). Comparison with the isotopic curve shows that the aeolian environment coincided with the most recent glacial periods of the Late Pleistocene. The relative sequence of deposition, as recorded in the K4 and K5 terraces, can be also applied to the K1–K3 terraces, with the shallow-marine deposits representing interglacial stages and the non-marine aeolian deposit representing a glacial stage. Each of the K1–K3 terraces comprises a basal marine deposit overlain by non-marine deposits, which are interpreted as representing a marine transgression, followed by a marine regression. The transgressive and regressive sedimentary sequence of each terrace is interpreted as representing a Pleistocene interglacial to glacial cycle. Extrapolating this concept allows the K1–K5 terraces to be tentatively correlated with several glacial cycles during the Late Pleistocene (Fig. 6). Based on this approach, the age of the K1–K3 terraces is suggested to be c. 0.5–0.2 Ma.
Several other age-related results have a bearing on any overall age model for the Pliocene–Pleistocene deposits (Fig. 6). OSL dating of coastal aeolianites from the Karpaz (Karpas) Peninsula, making up the eastern prolongation of the Kyrenia Range, has given ages of 1.5–0.4 ka, from the base to the top of the deposit (Erginal et al. 2012). These OSL dates indicate aeolian deposition during the most recent glacial period, consistent with our OSL results (Fig. 6). In addition, the marine gastropod Strombus bubonius has been identified within marine deposits of the K4 terrace near the modern coastline (Ferranti et al. 2006; Galili et al. 2015). Warm-water conditions associated with MIS 5e are thought to have resulted in proliferation of this gastropod throughout the Mediterranean Sea at this time (Galili et al. 2015). This, in turn, suggests that the marine deposits containing Strombus bubonius are likely to have accumulated from 130 to 114 ka (MIS 5), which is broadly consistent with our U-series dating (Fig. 6).
Prior to the Early Pliocene, the Kyrenia Range and the Troodos Massif had been tectonically juxtaposed in their present positions, separated by the Mesaoria seaway (McCallum & Robertson 1990, 1995b). Neotectonic faulting is known to have been active along the Ovgos (Dar Dere) Fault during the Plio-Pleistocene, to the south of the Kyrenia Range (Harrison et al. 2004; McCay & Robertson 2012b). The question arises as to whether the emergent Troodos and Kyrenia areas were uplifted at similar or different times.
The deposits on the southern side of the Mesaoria (Mesarya) Basin can be stratigraphically correlated with similar deposits on the northern side of the Mesaoria (Mesarya) Basin (McCallum & Robertson 1990). The presence of non-marine, fluvial deposits (Apalos Formation) around the periphery of the Troodos Massif indicates that this area was above sea level by the Late Pliocene (McCallum & Robertson 1995a; Schirmer et al. 2010; Kinnaird et al. 2011; Weber et al. 2011). Throughout the Mesaoria (Mesarya) Basin, the mudstone and marl facies of the Nicosia (Lefkoşa) Formation pass transitionally upwards into bioclastic calcarenites of the Athalassa (Gürpınar) Formation. Close to the Kyrenia Range the succession continues upwards into lagoonal to non-marine, deltaic facies (Palamakumbura 2015; Palamakumbura & Robertson 2016).
Based on the microfossil evidence from the southern part of Cyprus (e.g. Pissouri Basin), Lord et al. (2000) inferred an Early to Late Pliocene age for the Nicosia (Lefkoşa) Formation. At that time, the Pliocene–Pleistocene boundary was placed at 1.8 Ma, thus placing the Athalassa (Gürpınar) Formation at least partially within the Pliocene. However, the more recently redefined boundary at 2.6 Ma (Cohen & Gibbard 2010) places most of the Athalassa (Gürpınar) Formation within the Pleistocene, throughout Cyprus. The available facies and dating evidence, therefore, suggests that the Troodos Massif and the Kyrenia Range were uplifted more or less simultaneously (Fig. 7).
Rates of uplift
Uplift rates can be estimated based on the inferred ages and the heights of successive terraces, field-based correlations, and both geomorphological and sea-level correlations (Table 6). In calculating uplift rates, the height and age of adjacent terraces were taken into account. The heights of the K4 and K5 terraces are those of the quantitatively dated deposits. The heights of the K1–K3 terraces are taken from the marine facies at the base of each terrace, which is observed to be the oldest part of the terrace. No marine facies are preserved within the K0 terrace that could be indicative of a base level; however, a height of 1000 m is used to represent the base level of the K0 terrace, which marks the start of uplift. The ages of the K4 and K5 terraces were determined by U-series disequilibria and OSL dating, whereas the K1–K3 ages were determined by correlations with the Pleistocene marine isotopic stages (see Table 6).
Using the age information, the amount of uplift that took place between the accumulation of the marine deposits of two successive terraces can be calculated. This allows the amount of uplift to be calculated between equivalent shallow-marine deposits within two different terraces. The K0–K2 terrace transition was chosen because the K0 and K1 terraces represent contemporaneous depositional environments (Palamakumbura 2015; Palamakumbura & Robertson 2016); therefore, the K0–K2 calculated uplift rate allows the K0 terrace deposits to be compared with the marine base of the K2 terrace. The calculated K3–K4 terrace uplift rate has a large uncertainty owing to the marine base of the K3 terrace being poorly preserved. Using the above method, the uplift rate between the K0–K2, K2–K3, K3–K4 terraces, within the K4 terrace and also between the K4–K5 terraces can all be calculated (Table 6).
The highest uplift rates were for the K0–K2 terrace, of >2 mm a−1 during the Middle Pleistocene. Uplift rates begin to fall for the K2–K3 terrace to c. 1 mm a−1 during the Middle Pleistocene. Finally, during the Late Pleistocene (K4 and K5 terraces) uplift rates were reduced to c. 0.2 mm a−1. Two important implications are: first, the Kyrenia Range underwent rapid surface uplift during the Early to Middle Pleistocene and, second, the uplift was initially rapid and then slowed during the Middle to Late Pleistocene.
To put the calculated uplift rates in their regional context, it is interesting to draw comparisons with other areas of the easternmost Mediterranean that have also undergone major Pleistocene uplift (Table 7). First, areas to the east of Cyprus have undergone generally constant and slower uplift rates compared with the Kyrenia Range. The Lebanese coast, specifically, has apparently undergone constant, gradual uplift during the Pleistocene (Seyrek et al. 2008). Second, the calculated uplift rates for the Kyrenia Range and the Troodos Massif differ, the latter being generally faster. Nevertheless, the uplift rates of both the Troodos Massif and the Kyrenia Range appear to have been fastest during the Early Pleistocene, decreasing during the Middle and Late Pleistocene, as suggested by the age v. height AMSL model for both areas (i.e. for Troodos: Poole et al. 1990; for Kyrenia: this study).
Implication for processes of surface uplift
Cyprus and also southern Turkey (Schildgen et al. 2012) have undergone substantial Pleistocene uplift related to the convergence of the African and Eurasian plates. However, the scales of the uplifted areas vary regionally (Table 7), which needs to be taken into account in any tectonic interpretation. The narrow, uplifted Kyrenia Range is <6 km wide and >150 km long, whereas the domed uplift of the Troodos Massif is c. 50 km in circumference. The uplifted southern part of Anatolia and the Taurus Mountains is c. 150 km wide and >450 km long, whereas the roughly north–south-trending uplifted Amanos Mountains are c. 15 km across and c. 60 km long (Gomez et al. 2006). The uplifted Lebanese coastal area and is c. 200 km long and <50 km wide (Seyrek et al. 2008).
After a short-lived phase of thrusting, then emergence during Middle to Late Eocene time (Robertson et al. 2012), the Kyrenia Range became deeply submerged during the Oligocene to Mid-Miocene. This was followed by shelf-depth fine-grained deposition during the Late Miocene (McCay & Robertson 2012a), continuing after the Messinian salinity crisis during the Pliocene (Palamakumbura & Robertson 2016). In contrast, the Troodos Massif remained beneath a deep sea during Late Cretaceous to Oligocene time; initial uplift to shelf depths occurred during the Miocene, with localized emergence in the south prior to the Tortonian (Robertson 1977, 1990, 1998a). Following the Messinian salinity crisis, shelf-depth accumulation persisted around the periphery of the Troodos Massif until the Late Pliocene. Southern Anatolia was largely emergent by the Late Eocene following closure of a pre-existing Tethyan oceanic basin in the region (Clark & Robertson 2005). However, shallow-marine deposition persisted locally in southern Anatolia (e.g. in the Mut Basin) into the Pleistocene (Walsh-Kennedy et al. 2014). Southern Turkey was uplifted in stages over c. 23 myr, although uplift was fastest during the Pliocene–Pleistocene (Schildgen et al. 2012). In addition, the Lebanese coast and the Amanos Mountains regions were mainly uplifted during the Pliocene to Recent (Gomez et al. 2006; Seyrek et al. 2008).
Assuming the calculated uplift rates (Table 5) are broadly correct, the height differences between the Kyrenia Range and the Troodos Massif could simply reflect differences in uplift rates. However, assuming a similar uplift rate for the Kyrenia Range to that for the Troodos Massif, it is difficult to imagine the narrow Kyrenia Range ever reaching c. 2000 m AMSL, because it would have collapsed before reaching anything like this height. The lithologies of the Kyrenia Range are highly fractured and faulted related to successive Late Cretaceous, Mid-Eocene and Late Miocene–earliest Pliocene stages of deformation (Robertson & Kinnaird 2016). Such inherited rheological weaknesses would have favoured rapid disintegration of the spine of the Kyrenia Range during and soon after rapid uplift.
Even at its present height the oldest terrace (K0, equivalent to the Karka terrace; Ducloz 1972; Dreghorn 1978) is dominated by megabreccias with blocks up to tens of metres in size, which document just such a collapse process (Palamakumbura 2015; Palamakumbura & Robertson 2016). No such sediments are preserved around the more broadly uplifted Troodos Massif (Poole & Robertson 1991).
Taking account of the extrapolated age data (see Table 5) and other constraints (e.g. lithology; structure), the Kyrenia Range is estimated to have been uplifted during the Early Pleistocene at >2 mm a−1, which is significantly faster than for adjacent areas of the easternmost Mediterranean region. The Troodos Massif appears to have been uplifted at a maximum rate of 2.4 mm a−1, mostly during the Early Pleistocene (Poole et al. 1990). The massif was uplifted as a broad dome, supported by erosionally resistant ophiolitic diabase. In contrast to the Kyrenia Range, mass wasting (as opposed to normal fluvial deposition) is seen only locally; for example, around parts of the serpentinite core where the topography is unusually steep (e.g. Amiandos area; Poole & Robertson 1998).
Despite the differences in inferred maximum uplift rate and lithology between the Kyrenia Range and the Troodos Massif it appears that the timing and rates of uplift of the two areas are more similar to each other than to those of any of the adjacent areas discussed above (e.g. southern Turkey; Amanos Mountains, Lebanese coast). This, in turn, points to a single overall driving mechanism for the uplift of both the Troodos Massif and the Kyrenia Range. The uplift of the Troodos Massif has previously been explained by the collision of the Eratosthenes Seamount, a crustal edifice, with the Cyprus trench to the south of Cyprus (Robertson 1990, 1998c, 2000; Robertson & Grasso 1995; Kempler 1998; Mart et al. 1997). It has also been suggested that the domal Troodos uplift might relate to sepentinization of the underlying mantle wedge (Robertson 1990, 2000; Poole & Robertson 1991; see also Shelton 1993). The collision of the Eratosthenes Seamount could also have driven the uplift of the Kyrenia Range, assuming that this was kinematically coupled with the Troodos Massif, explaining the similarities in the timing and rates of uplift of the two areas (Fig. 7). The entire process occurred within a region of dominantly compressional tectonics (Le Pichon & Kreemer 2010; Seyrek et al. 2014), which was locally exaggerated by the collision of the Eratosthenes Seamount with the Cyprus active margin. This would also explain why the Kyrenia Range reaches its maximum height at a similar longitude to both the Eratosthenes Seamount and the Troodos Massif, and also why the Kyrenia Range submerges both to the east and the west away from the inferred north–south collision zone.
Several alternative driving mechanisms for the regional-scale uplift of the Troodos Massif and the Kyrenia Range can also be considered. One involves strike-slip or transpression during the Pliocene–Pleistocene tectonic escape of Anatolia towards the Aegean region (Harrison et al. 2004; Jaffey & Robertson 2005). Strike-slip has indeed been found to play an important role related to the Ovgos (Dar Dere) Fault, which essentially separates crust related to the Kyrenia Range from that related to the Troodos Massif (Harrison et al. 2004; McCay & Robertson 2012b). However, there is no structural evidence of transpression of Pliocene age either within the Kyrenia Range (Palamakumbura 2015) or from scarce published seismic reflection data (Aksu et al. 1992; Calon et al. 2005a). Another alternative involves crustal extension. Extensional tectonics have played an important role in the easternmost Mediterranean region (Glover & Robertson 2004). Offshore seismic data from the Cilicia and Latakia Basins adjacent to Cyprus indicate the presence of extensional structures within pre-Pliocene units based on offshore seismic data (e.g. Calon et al. 2005b). Onshore field studies in southern Turkey are also indicative of crustal extension, as noted around the periphery of the Gulf of Iskenderun (Boulton & Robertson 2008). In addition, the Hatay Graben is inferred to represent a transition from extension to transtension during the Miocene to Pleistocene, probably as a consequence of the regional westward tectonic escape of Anatolia (Boulton & Robertson 2008). Also, to the west of the Eratosthenes Seamount, crustal extension took place within the Polis Graben of western Cyprus, continuing into the Pleistocene (Payne & Robertson 1995). Such areas affected by Pleistocene extension are located well to the east or the west of the inferred roughly north–south Eratosthenes–Troodos–Kyrenia collision zone.
The Plio-Pleistocene uplift of the southern Anatolian plateau and the Taurus Mountains generally has recently been attributed to the break-off, tearing or delamination of the down-going African plate in the region (Faccenna et al. 2006; Cosentino et al. 2012; Schildgen et al. 2012). Such slab disruption might also have influenced the uplift of the Kyrenia Range. However, the Kyrenia Range is located 200 km south of the inferred location of the tear in the downgoing slab, whereas the Anatolian Plateau is located directly above this inferred feature (Schildgen et al. 2014).
The rapid uplift rates inferred during the Pleistocene uplift of the Kyrenia Range suggest that tectonic uplift was the dominant control on deposition during this time, modified by the effects of sea-level change and related climatic change. The uplift rates and preserved altitudes of the older and higher terraces are much greater than could be explained by eustatic sea-level change on its own. However, the inference that tectonic uplift slowed markedly during the Middle to Late Pleistocene means that eustatic sea-level change is likely to have played a relatively greater role in the formation of the younger and lower terraces.
The uplift of the Kyrenia Range occurred within a regional-scale plate convergence zone in the Eastern Mediterranean region and can thus provide insights into fundamental crustal processes.
Shallow-marine sediments, mostly open-marine marls that predated major uplift, are dated as Pliocene, using a combination of available microfossil dating and mostly existing Sr isotopic evidence.
New palaeomagnetic evidence for overlying shallow-marine carbonate sediments indicates accumulation during reversed and normal polarity intervals, consistent with an Early Pleistocene age.
U-series dating of aragonitic Cladocora coral from the youngest of four uplifted littoral marine terraces (K4 terrace) along the northern coast of Cyprus indicates deposition at >350, 243, 131 and 127 ka (equivalent to MIS 9, 7 and 5).
OSL dating of aeolian grainstones overlying the youngest marine terrace (K5) indicates accumulation from 76 to 53 ka.
Correlation of the available age data (new and existing) from the littoral-marine terrace deposits with the eustatic sea-level curve based on the oceanic oxygen isotope record indicates that the dated coral-bearing littoral deposits accumulated during eustatic sea-level highs, whereas the dated aeolianites formed during regressive periods.
The well-established dates for the younger (K4 and K5) littoral terraces can be used to calculate rates of uplift of <0.2 mm a−1 during the Late Pleistocene.
The older, topographically higher littoral terraces along the northern flank of the Kyrenia Range can be correlated with earlier major marine transgressions (MIS 13, 11 and 9), suggesting that the Kyrenia Range was uplifted at maximum rates of either 1.2 or >2 mm a−1 during early Pleistocene time.
Comparison of the inferred uplift rates of the Kyrenia Range and the Troodos Massif with those for adjacent areas of the easternmost Mediterranean region suggests similar timings and rates of uplift of the two crustal units in Cyprus, which differ from those for southern Turkey and the Lebanese coast.
The Pleistocene uplift of the Troodos Massif can be explained by the collision of the Eratosthenes Seamount with the Cyprus trench, and the near-simultaneous uplift of the Kyrenia Range directly to the north is likely to have a similar cause.
Acknowledgements and Funding
R.N.P. acknowledges the NERC CASE scholarship at the University of Edinburgh. Additional field and laboratory work was aided by the DARIUS Programme to A.H.F.R. and T.C.K. We are grateful for the additional financial support provided by the John Dixon Memorial Fund. R.N.P. also thanks É. Deady for help with the palaeomagnetic sampling. We thank M. Necdet for assistance with logistical aspects during the fieldwork. We also thank C. Mac Niocaill (University of Oxford) and C. Xuan (University of Southampton) for the use of palaeomagnetic laboratories and guidance with laboratory methods. A. Morris helped with interpretation of the palaeomagnetism data. We thank N. Odling for carrying out the XRD determination on the coral samples. A. Kelly kindly carried out the strontium dating at SUERC, East Kilbride. In addition, we thank R. Ellam and D. Sanderson for the use of the ICP-MS laboratory and the luminescence laboratories respectively, both at SUERC. The paper benefited from detailed comments from two anonymous referees.
- Received October 28, 2015.
- Revision received February 20, 2016.
- Accepted March 22, 2016.
- © 2016 The Author(s)
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