The tectonostratigraphic evolution of Cenozoic basins of the
Northern Tethys: The Northern margin of the Levant Basin
, Remy Deschamps
, Vasilis Symeou
, Christine Souque
, Christian Gorini
Fadi Henri Nader
, and Christian Blanpied
Geosciences Division, IFP Energies nouvelles, 1-4, avenue de Bois-Pre
´au, 92852 Rueil-Malmaison Cedex, France
ISTeP, UMR 7193, CNRS, 75005 Paris, France
Received: 30 April 2018 / Accepted: 5 November 2018
Abstract. The easternmost part of the Mediterranean corresponds to a tectonically complex region which is
linked with the convergence between Africa and Eurasia. The tectonostratigraphic evolution of this region is
poorly constrained because of the absence of exploration wells. Cyprus is a crucial area to assess the link
between the tectonic deformation and the consequent sedimentation in the Northern Levant margin. Paleogene
and Neogene basins in the southern part of Cyprus record the main tectonic events related to the convergence of
Africa and Eurasia. The objective of this contribution is to investigate the timing and the mechanisms of basin
deformation, as well as the sedimentary inﬁll of basins located onshore Cyprus and ﬁnally resolve how their
evolution is linked to the regional geodynamic events. Based on ﬁeldwork studies we reconstructed the
tectono-stratigraphic evolution of the Polis Basin and the Limassol Basin to propose a conceptual model for
the evolution of the Northern Levant margin, in accordance with the main geodynamic events. It is expected
that analysis of the Polis and Limassol depressions, and later comparison of them will also shed more lights on
the impact of the substratum and how it is associated to the main tectonic events.
Cyprus is located in the eastern corner of the Mediterranean
Sea. It is 225 km long (E-W) and 95 km wide (N-S) and is
bounded by the Cyprus Arc and Eratosthenes Seamount to
the south and Turkey to the north (Fig. 1). Eratosthenes
Seamount is the bathymetric expression of an isolated
carbonate platform which appears to subduct beneath the
Cyprus Arc (Papadimitriou et al., 2018;Robertson et al.,
2012). The Cyprus Arc is a complex structure which records
the opposite movement of African and Eurasian plates that
extends from the Ionian islands of western Greece to
Turkey, striking E-W (Glover and Robertson, 1998;
Kinnaird, 2008;Robertson et al., 2009).
The opening and the closing of Neotethys are recorded
onshore Cyprus by the formation and the subsequent
obduction of the Troodos Ophiolites during the Late
Cretaceous (Garfunkel, 1998,2004;Robertson and
Xenophontos, 1993;Stampﬂi and Borel, 2002). After the
obduction of the Ophiolites, a series of smaller basins were
formed (Kinnaird, 2008). These basins are: (a) Mesaoria
Basin; (b) Psematismenos Basin (PSB), (c) Limassol Basin
(LB), and (d) Polis Basin (PB) (Fig. 2A).
Different scenarios attempt to describe the tectonic
evolution of the region after the obduction of Troodos
Ophiolites. These scenarios referred to as (a) the advance
collision model which is associated with thrusting and
folding onshore Cyprus since the late Eocene (Calon
et al., 2005a, b); (b) the subduction and incipient collision
scenario that suggests a northward dipping subduction zone
to the south of Cyprus during the late Oligocene to early
Miocene and a recent collision of the Cyprus Arc with
Eratosthenes Seamount (Robertson et al., 2012);
(c) strike-slip scenario which depicts that the obduction of
the Troodos Ophiolites and the formation of the Neogene
basins is controlled by sinistral strike-slip (Harrison,
2008); (d) thrust belt forward propagation model which
suggests successive compressional pulses since the early
Miocene and a change to strike-slip in the early Pliocene
and/or Early Pleistocene due to the westward movement
of the Anatolia microplate (Reiche et al., 2016;Robertson
et al., 2012;Symeou et al., 2017).
This study is focused on the impact of tectonics on
sedimentation during the Miocene and is based on sedimen-
tological investigations that were undertaken during two
ﬁeld campaigns in the Polis and Limassol basins
(Fig. 2A). The Limassol and Polis basins are bounded by
a transverse zone which extends from the Kouklia village
to the Xeropotamos River (Fig. 2A). Monnet (2005) has
Dynamics of sedimentary basins and underlying lithosphere at plate boundaries: The Eastern Mediterranean
F.H. Nader, R. Littke, L. Matenco and P. Papanastasiou (Guest editors)
* Corresponding author: email@example.com
This is an Open Access article distributed under the terms of the Creative Commons Attribution License (http://creativecommons.org/licenses/by/4.0),
which permits unrestricted use, distribution, and reproduction in any medium, provided the original work is properly cited.
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REGULAR ARTICLEREGULAR ARTICLE
shown that this zone has a sigmoid geometry due to the
existence of a sinistral component on two main groups of
thrusts located to the north and the south of the Mamonia
Suture Zone (Fig. 1). It is expected that through basis
analysis (Polis and Limassol basins), will also shed light
on the impact of the substratum how it is linked with the
main tectonic events that are recorded onshore Cyprus.
2 Geological framework
Recent tectonic reconstructions of the Eastern
Mediterranean illustrated that different tectonic events
inﬂuenced the island of Cyprus (Harrison et al., 2004;
Kinnaird, 2008). The ﬁrst episode occurred during the Late
Cretaceous with the subduction of the African Plate
beneath the Eurasian Plate and subsequent obduction of
the Troodos Ophiolites (Dilek and Sandvol, 2009;Hawie
et al., 2013;Kinnaird, 2008;Robertson et al., 2009,2012).
The emplacement of the Troodos Ophiolites was
succeeded by their juxtaposition with the Mamonia
Complex (Bailey et al., 2000). The mechanism of the
juxtaposition of the two complexes – Troodos Ophiolites
and Mamonia Complexes – is still debated, and several
models have been proposed (Bailey et al., 2000;Lapierre
et al., 2007;Malpas et al., 1992,1993;Robertson and
Woodcock, 1979). Paleontological results of sediments that
pre- and post date this tectonic event indicate that the
juxtaposition ended during the Maastrichtian (Bailey
et al., 2000).
From the late Maastrichtian until the late Eocene,
Cyprus was covered by deep marine sediments rich in
planktonic foraminifera and calcareous nannofossils, which
correspond to the ﬁrst three members of the Lefkara
¨hler and Stow, 1998;Kinnaird, 2008;
Robertson, 1977;Fig. 3). The Lefkara Formation is subdi-
vided into four (4) geological units: (a) the Lower Marl Unit
consists of pinkish marls, (b) the Chalk and Chert Unit,
(c) the Chalk Unit, and (d) the Upper Marl Member which
is composed of marly chalks (Ka
¨hler and Stow, 1998;
Fig. 3). Robertson (1977) identiﬁed some in-situ benthic
foraminifera in the upper levels of this unit and proposed
a gradual shallowing of the area during the late stages of
Eocene (Fig. 3).
The transition between the Lefkara (Paleogene) and
the overlying Pakhna (Miocene) formations is highly
connected the collision between the African and the
Eurasian plates (Dercourt et al., 1986;Jolivet and
Faccenna, 2000). In some places, there is a gradual change
in the depositional environment, whereas in others the
transition is marked by erosional surfaces (Follows, 1992,
1996). Pakhna Formation shows a lateral change in facies
from reefs to hemipelagit and turbiditic deposits (Eaton
and Robertson, 1993;Payne and Robertson, 1995;
Fig. 3). The reefs referred to as the Terra and the Koronia
members and can be found at the base and the top of the
Miocene succession (BouDagher-Fadel and Lord, 2006;
Follows, 1992,1996;Fig. 3). Early Miocene reefs grew as
upstanding patches under a relatively deep and calm sea,
on isolated carbonate shelves (BouDagher-Fadel and Lord,
2006;Follows, 1992) and consist of several coral frame-
stones (e.g., faviids sp.andPorites sp.). In contrast, the
Koronia Member is a bindstone that is comprised of
monospeciﬁc, laminar poritid corals (Follows, 1996)and
Fig. 1. Bathymetric map of the Eastern Mediterranean showing the topography as well as the most important structural elements
bounding the Levant Basin (modiﬁed from Hawie et al., 2013).
N. Papadimitriou et al.: Oil & Gas Science and Technology - Rev. IFP Energies nouvelles 73, 77 (2018)2
Fig. 2. [A] Simpliﬁed geological map of Cyprus illustrating the distribution of the basement terranes: the Troodos Ophiolites in
purple, the Mamonia Complex in brown, the Keryneia Range in grey – and sedimentary cover and the main sedimentary basins.
A distinction is made between circum-Troodos and circum-Keryneia sediments. Abbreviations as follows: AL – Akrotiri Lineament;
YFS – Yerasa Fault System; OFS – Ovgos Fault System; ATFS – Arakapas Transform Fault; PFS – Pafos Fault System; LB –
Limassol Basin; PIB – Pissouri Basin; PSB – Psematismenos Basin; PB – Polemi Basin (modiﬁed from Kinnaird, 2008); [B] Overview
of the dataset used for the present study.
N. Papadimitriou et al.: Oil & Gas Science and Technology - Rev. IFP Energies nouvelles 73, 77 (2018) 3
was developed in a shallower sea with varying turbulence
Pakhna Formation is overlain by Messinian deposits
(Kalavasos Formation) which are exposed in localized
depocenters such as Polemi, Pissouri, and Psematismenos,
and they consist of alternating sapropels, marls and
carbonates (Manzi et al., 2016;Figs. 2 and 3). Finally, the
evaporitic sequence is succeeded by the calcarenites and
the relatively deep-water marls of the Nicosia Formation,
which in turn passes into stacked ﬂuvial deposits that con-
sist of well-bedded gravels, ﬂood deposits, and palaeosols
(Kinnaird, 2008;Figs. 2 and 3).
Fig. 3. A synthesized chronostratigraphic chart with the main lithologies onshore Cyprus (modiﬁed from Kinnaird, 2008).
N. Papadimitriou et al.: Oil & Gas Science and Technology - Rev. IFP Energies nouvelles 73, 77 (2018)4
In order to propose a conceptual model which illustrates the
major geodynamic events and their impact on sedimentary
processes, ﬁeldwork in the southern part of Cyprus was
The ﬁeld campaigns were focused on the description of
the lithology and the stratigraphic contacts between the
Paleogene and Miocene formations (Figs. 2 and 4). During
the ﬁeldwork six (6) composite sedimentary logs were
completed in the two areas of interest and sixty-ﬁve (65)
sediment samples were recovered (Fig. 2B). The locations
of logging were thoroughly chosen based on accessibility,
dating requirements (to constrain the Upper Cretaceous
to Pliocene rock successions), and variety of the facies
attested during the Miocene. Thirty-ﬁve (35) samples were
used for dating to have a better time constraint regarding
the continuity of stratigraphic contacts and the sedimen-
tary facies extent (Fig. 2B). In particular nannofossil
analyses involved counting all specimens within a standard
traverse (60 ﬁelds of view at ·1000 magniﬁcation) as well as
scanning of the rest of the slide for rare species. Each of the
samples is dated using only reﬁned versions of the
‘‘standard’’ published biostratigraphic zonation schemes
(Martini, 1971;Sissingh, 1977,1978). Moreover, petrogra-
phy studies have been conducted to examine the microtex-
tural features of the studied rock interval. For petrographic
studies, sixty-ﬁve (65) thin sections were prepared and
studied with a polarizing microscope at 1250 magniﬁca-
tions, following the Dunham’s classiﬁcation. Also, the ﬁeld-
work was supported by hydrogeological, exploration and
geotechnical boreholes which are drilled by the Geological
Survey Department of Cyprus (GSD) and provide subsur-
face data on borehole chips or boreholes cores. These
drillings resulted in borehole data ranging up to 300 m that
are used here in order to document the basement of the
Polis Basin. Out of the provided data (courtesy of GSD),
four boreholes were chosen and analyzed for the purposes
of this study (Fig. 4).
Five (5) facies associations are identiﬁed onshore Cyprus
from the Late Cretaceous until the Pliocene period. Each
of the deﬁned facies allowed us to deﬁne the characteristics
for the formations recognized onshore Cyprus and portray
the architecture of Polis and Limassol basins since the
4.1 Facies analysis
4.1.1 Intertidal deposits
Facies association 1
Description. Fa 1A is characterized by an alternation of
clays (0.2 m) with cross-bedded (0.2 m), bioclastic, dolomi-
tized grainstones (Fig. 5.1). The laminated clays show
traces of mud-cracks and contain ﬂat pebbles (Fig. 5.1B),
while at the base of the section, the grainstones show traces
of mega ripples (Fig. 5.1D).
Interpretation. This facies refers tidal-ﬂat deposits
which are mainly found in the intertidal zone. Exposures
of this facies are mainly located along the northern parts
of Limassol Basin (3443019.7600N; 33102.1900E).
Description. Fa 1B consists of massive, parallel beds of
coarse-grained, bioclastic grainstones, alternating with
bioturbated wackestones rich in benthic foraminifera
(Figs. 5.2B and C). Bioturbation is occasionally identiﬁed
and corresponds to Thalassinoides burrows (Eaton and
Interpretation. The grain textures of these rock units, as
well as the traces of Thalassinoides, suggest a shallow, high-
energy environment and more particular it represents sand
shoal deposits (Table 1). The bioclastic material was mostly
derived from patch reefs. In contrast, the allochthonous
calcarenites have been transported from the Troodos
Ophiolites and/or the Mamonia Complex (Follows, 1992,
1996;Eaton and Robertson, 1993). Exposures of this facies
are mainly located along the northern parts of Polis
(3447016.0700N; 3254053.9900E) and Limassol basins
4.1.2 Barrier reef/open shelf deposits
Facies association 2
Description. Fa 2A corresponds to rudstone/boundstone
interval with corals, red algae and bivalves (Fig. 2A)associ-
ated with bioclastic packstone (red algae and bivalve deb-
ris). The primary frame builders of this unit are the
Porites sp.(Fig. 6.1B) which in some cases are abundant
while in others their number decreases towards the top of
the sequence, and they are replaced by Poritis sp. Secondary
frame builders found in grainstone and rudstone textures
(Fig. 6.1C) include red algae and small benthic foraminifera
(miliolids, nummulitids, echinoderms, bivalves).
Interpretation. The presence of the in-situ corals
(i.e., Porites sp.) reveals that these sediments deposited
in a very shallow and protected euphotic environment
(Fig. 6.1B and Table 1). Usually, this facies has been recog-
nized in circular shape structures attaining a width of
100 m and a thickness of 80 m and has been described to
Description. Fa 2B facies presents grainstone grading into
wackestones rich in red algae (Fig. 6.2B ) benthic forami-
nifera (Fig. 6.2B,i.e.,  crinoids and  Lepidocyclina)and
coral debris. Incisions of structureless rudstone units (1–2 m
thick), with abundant bioclasts, bioturbation, benthic fora-
minifera (Fig. 6.2C;  miliolids) appear to cut this wacke-
stone to grainstones beds.
Interpretation. This facies indicates a high-energy envi-
ronment and the abundant red algae and coral fragments
suggest a reef front setting (Fig. 6.2). Reef front deposits
have been observed in a small distance from in-situ reefs
of early Miocene age (i.e., Burdigalian, Terra Member),
mainly in the Polis Basin (3453027.3500N, 3221056.7700E).
N. Papadimitriou et al.: Oil & Gas Science and Technology - Rev. IFP Energies nouvelles 73, 77 (2018) 5
Fig. 4. Simpliﬁed geological map of the western part of Cyprus showing the main geologic formations with their ages (adapted from
Cyprus Geological Survey Department, 1995).
N. Papadimitriou et al.: Oil & Gas Science and Technology - Rev. IFP Energies nouvelles 73, 77 (2018)6
Fig. 5.  (A) Studied outcrop near Koilani village (3450015.5200N; 3251031.1200E); (B) Thin bedded bioclastic grainstones;
(C) Photomicrographs (PPL – Plane Polarized) of packstone/grainstones with abundant benthic and pelagic foraminifera; (D) Traces
of mega-ripples in a bioclastic grainstone;  (A) Studied outcrop of the upper Pakhna Formation in the eastern part of Limassol
Basin (3443019.7600N; 33102.1900E); (B) Grainstones interbedded with wackestones; the base of the calcarenitic unit is composed of
allochthonous grains; (C) Bioturbated w/p grading to grainstone; (D) Photomicrograph (PPL – Plane Polarized) of calcarenite with
shell debris and quartz.
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Facies association 3
Description. Fa 3A consists of highly-bioturbated pack-
stones, intercalated with less bioturbated wackestones
(Fig. 6.3). Both packstones and wackestones are rich in
benthic ( crinoids  Lepidocyclina) and planktonic fora-
minifera ( Globigerina;Fig. 6.3C)aswellasshelldebris
(Fig. 6.3). Red algae debris is only occasionally present
(Fig. 6.3D), whereas coral fragments are absent. The vari-
ety of fauna in these textures (bivalves, and benthic forami-
nifera, such as (Fig. 6.3D  Lepidocyclina and  crinoids),
indicate normal salinities and, thus, open-shelf conditions
Interpretation. This facies is interpreted as open shelf
deposits which are mainly exposed in the Pegeia region
(i.e., in the western side of the Polis Basin 3453010.5700N,
4.1.3 Slope-deep marine deposits
Facies association 4
Description. Fa 4A is represented by alternations of biotur-
bated marlstone with packstones (Figs. 7.1A and B). This
unit is highly-bioturbated, with abundant foraminifera
(benthic and planktonic) and shell debris (Fig. 7.1C). The
packstones are cross-bedded and have sharp contacts with
the laminated marls (Figs. 7.1A and B).
Interpretation. The erosional surface at the base of the
grainstones and their gradual upward ﬁning suggest that
this sedimentary facies corresponds to proximal slope
turbiditic deposits. However, the absence of water escape-
structures and sole-marks, suggest that these sediments
correspond to calciturbidites (Table 1). Previous works
have deﬁned that these sediments are Pakhna Formation
equivalents and are mainly exposed in the Limassol Basin
(Eaton and Robertson, 1993;Kinnaird, 2008).
Description. Fa 4B consists of bioclastic grainstones, which
ﬁne upwards, and are intercalated with laminated marls
(Figs. 7.2A and B). The whole unit is highly-bioturbated
and contains abundant planktonic foraminifera (Fig. 7.2A)
and is found in the small cyclic sequences (Fig. 7.2C). More-
over, the grainstones are cross-bedded, while the sedimen-
tary structures faint upwards.
Interpretation. The description given for these deposits,
seem to agrees with the description that was proposed by
Stow and Lovell (1979) and Stow et al., (2002) for con-
tourite systems (Table 1). Ka
¨hler and Stow (1998) have
shown strong evidence for deposition of contourites from
the early (eastern part of the island) and late Oligocene
(western part of the island) until the early Miocene.
Table 1. Summary of the facies associations’ description and proposed depositional environments for the Upper
Cretaceous to Upper Miocene rock interval of southern Cyprus.
Facies associations Facies description Depositional environment
Restricted shallow marine Fa 1 A Early dolomitized bioclastic grainstone with
current ripples and mega-ripples alternated by
bioturbated wackestone with some bioclasts
and mud cracks.
B Alternation of bioclastic grainstone rich in
miliolids and some lepydocyclina with
bioturbated bioclastic wackestone to packstone.
Barrier reef Fa 2 A Boundstone/bafﬂestone with red algae debris,
corals in place, Lepidocyclina and bioclasts. Reef /Back-reef
B Breccia (boundstone with red algae debris,
coral debris, large benthic foraminifera
(Lepidocyclina, bryozoans) occasional incisions.
Open shelf deposits Fa 3 A Wackstones to packstones with an abundance
of benthic foraminifera, and bioclasts. Erosional
contact between the packstone and the
Slope deposits Fa 4 A Intercalation of grainstone packstone beds
(highly bioturbated) with marls. Slope Deposits/ turbidites
B Large contorted fragments of laminated marls
and carbonates 3, 5 m thick. MTCs
C Bioclastic grainstones which are ﬁning upwards
and intercalated with a marly laminated unit
including benthic foraminifera.
Toe of the upper slope
Basin Fa 5 A Mudstone to wackestone with bioclasts, pelagic,
traces of zoophycos. Basinal setting
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Fig. 6.  (A) Reef limestones found near the Androlykou quarry (35000.9100N; 322307.7000E); (B) Photograph showing corals that
are probably in place; (C) Photograph showing coral fragments in a boundstone;  (A) Photograph showing a studied outcrop of reef
front deposits in the Pegeia region (3453027.3500N, 3221056.7700E); (B) Photomicrograph (PPL – Plane Polarized) of a packstone to
grainstone with benthic foraminifera ( red algae cross-section;  crinoids;  Lepidocyclina;  miliolids) (C) Photomicrographs
(PPL – Plane Polarized) of (C) Packstone with medium to low porosity;  (A) Studied outcrop of open shelf deposits (Fa 3A) in the
Pegeia region (3453010.5700N; 2320052.200E); (B) Wackestone with some bioturbation; (C, D) Photomicrographs (PPL – Plane
Polarized) of (C) Packstone with medium to low porosity ( crinoids;  Lepidocyclina;  Globigerina planktonic foraminifera);
(D) Packstone with benthic foraminifera ( Lepidocyclina;  miliolids;  red algae).
N. Papadimitriou et al.: Oil & Gas Science and Technology - Rev. IFP Energies nouvelles 73, 77 (2018) 9
Fig. 7.  (A) Photograph showing a turbidite system in the Limassol Basin (3448023.3000N; 3247029.4900E); (B) Wackestone
grading to packstone; (C) Highly bioturbated mudstone grading to packstone;  (A) Micritic non-porous mudstone, rich in pelagic
foraminifera; (B) Cross-bedded grainstones ﬁning to marlstones; (C) Mudstone overlain by gross bedded grainstones suggesting the
initiation of a new cycle;  (A) Panoramic view of the MTCs identiﬁed along the Limassol-Pafos highway (3445021.3000N;
3445021.3000E);  (A) Studied outcrop of basinal sediments near Monagri village (34470016.0700N; 3254053.9900E); (B) Alternations of
mudstone/wackestones with packstones; (C) Highly bioturbated mudstone; (D) Photomicrograph (PPL – Plane Polarized) of
wackestone with benthic and planktonic foraminifera.
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Description. Fa 4C is described by chaotic intervals, usually
3.5 m thick, which are composed of large, contorted
fragments of laminated marls (Fig. 7.3). These chaotic
intervals, pinch out on both sides and are characterized
by a basal shear zone (30–150 m long), as well as upward
curved or stepped margins, indicating several distinct lobes.
Interpretation. The Fa 4C has been interpreted as Mass
Transport Complexes (MTCs; Eaton and Robertson, 1993;
Lee and Stow (2007);Lord et al., 2009). The MTCs have
mainly been identiﬁed in the western and eastern sides of
the Limassol Basin. In the western part of the Limassol
Basin the MTCs are found along the recently constructed
Limassol-Pafos highway (Fig. 7.3;3441034.6600N;
3251053.4700E) as well as in an outcrop 4 km north from
Anogyra village (3445013.7200N; 3241012.9900E). Biostratig-
raphy results from Lee and Stow (2007),Lord et al., (2009)
and Kinnaird (2008) have shown that the MTCs are
interbedded with late Miocene sediments.
Description. This facies consists of thin, alternating beds of
wackestone and marls, interbedded with packstones
(Fig. 7.4A). All of the packstones and wackestones are
highly bioturbated, with benthic and planktonic foramini-
fera and bioclasts (Figs. 7.4B–D).
Interpretation. The bioturbation found in these
textures is characterized by Zoophycos,Chondrites and
Planolites suggesting a deep-water setting (e.g., open shelf;
700–1100 m; Eaton and Robertson, 1993). This facies is
exposed in several outcrops in Limassol Basin and marks
the most common appearance of Pakhna equivalent. These
outcrops are also described by Eaton and Robertson, (1993)
and are referred to as Basin-plain Association. In particular,
they described these sediments as off-white with a high
planktonic/benthic foraminifera ratio (i.e., 85–95% Eaton,
1987;Eaton and Robertson, 1993).
4.2 Architecture of the basins in southern Cyprus
Through the sedimentological, the structural observations
as well as spatial distribution of the facies association we
propose a new tectonostratigraphic framework of these
two basins and attest the different controlling factors (such
as tectonics, eustatism, and paleoenvironmental condi-
tions), which contribute into the geological evolution of
the southern part of Cyprus.
4.2.1 Polis Basin
The Polis Basin is located in the westernmost part of
Cyprus (Figs. 2 and 4). To the west, the topography of
the basin (i.e., the Structural unit 1 – Pegeia region) is
controlled by N-S trending thrusts (Fig. 4). Similarly, to
the north (i.e., Structural unit 2-Akamas Peninsula), the
late Miocene reefs are occasionally sitting on top of the
Mesozoic sediments (Fig. 4). The largest Structural unit 3
in the Polis Basin is referred to as ‘‘Polis Graben’’ and
corresponds to an asymmetrical depression NNW-SSE
oriented that is approximately 20 km at its widest point
(Fig. 4). Finally, the Structural unit 4 is located along the
eastern ﬂank of Polis Graben along NW-SSE trending
faults (Fig. 4).
Structural unit 1 (Pegeia region)
The Pegeia region is bounded to the northeast by the
Kathikas Formation which is juxtaposed to the Quaternary
marine terraces (Fig. 4). The boundary between the Lefkara
and Pakhna formations found in this region is uncon-
formable (Fig. 8.1). For instance, the Late Cretaceous
sediments are overlain by the Miocene Formation while
the Paleogene succession (Lefkara Formation), is locally
absent (Figs. 8.1A–C). This contact also appears to be
diachronous. In particular, towards the coastline, the early
Miocene sediments rest unconformably on top of the
Lefkara Formation (Fig. 8.1)whiletowardstheeast
(i.e., Kathikas village; Fig. 4) the Lefkara Formation
pinches on the slopes of the Kathikas Formation (Fig. 4).
Swarbrick and Naylor, 1980 observed that the Cretaceous is
(lowest part of the composite log in Fig. 8.1A)iscomposed
of poorly sorted brownish to grey clasts supported by
argillaceous matrix. The beds of this interval are 1–3 m
thick, and they can be deﬁned by the variation in the size
of clasts and fabrics, or the interbeds of pelagic chalks
(Fig. 8.1A). Coccolith fragments and foraminifera found
within the pelagic chalks dated back to late Maastrichtian
indicating pauses in debris ﬂow sedimentation (Urquhart
and Banner, 1994;Morse, 1996). We propose that these
sediments are deposited in a deep marine bathyal setting
as a result of gravitational forces.
The Cretaceous interval is overlain by a sequence which
consists of alternations of wackestone to mudstone beds,
with pelagic foraminifera (40–120 m, Fig. 8.1A) suggesting
basinal deposits (Fa 5A). Biostratigraphy results have
shown that these sediments were deposited during the early
The Eocene–Miocene stratigraphic contact in the Pegeia
region is occasionally erosive (Figs. 8.1B and C). The
overlying unit (120–150 m, Fig. 8.1A) corresponds to reef
front deposits (Fa 2B). This facies is occasionally cut off by
channelized conglomerate sheets (1–2 m thick), with an
abundance of red algae fragments and bioclasts and
is found in several locations (near Kathikas Village:
3453045.5800N, 322306.7400E; Pegeia village: 3453027.7000N;
3221016.4600E). The presence of the calcareous nannofos-
sils such as Sphenolithus heteromorphus (restricted to
Nannofossil Zones NN5-NN4), Helicosphaera ampliaperta,
and H. scissura conﬁrms an early Miocene age (Burdigalian;
In-situ early Miocene reefs (Fa 2A; Table 1) have been
identiﬁed along the south-western coastline of Cyprus
and near Androlykou village (Fig. 4). The majority of these
reefs appear to be 50 m thick, with few signs of deformation
N. Papadimitriou et al.: Oil & Gas Science and Technology - Rev. IFP Energies nouvelles 73, 77 (2018) 11
Fig. 8.  (A) Sedimentary log measured in the Pegeia region showing the main sedimentary facies from the Late Cretaceous to late
Miocene as well as the ages subdivisions deduced from biostratigraphic analysis; (B) Panoramic view of the contact between the
Miocene (Pakhna Fm), Paleogene (Lefkara Fm) and Late Cretaceous (Kathikas Fm) sediments. The contact between the Miocene
and the Cretaceous sediments is shown by the red line (3455033.8100N; 3224057.0900E); (C) Karstiﬁcation identiﬁed near Kathikas
village.  (A) Synthetic log showing the main lithologies (interpreted during the present study) recovered from of the Borehole 1980/
076; (B) Messinian evaporites on top of Miocene neritic carbonates in the southern part of Polis Graben near Theletra village
(3454036.4000N; 3227032.2700E).  (A) Sedimentary log illustrating the main facies and the depositional environments in the eastern
ﬂank of Polis Basin; Panoramic view of the Koronia Member sitting on top of Late Cretaceous sediments in the north-eastern ﬂank of
Polis Basin (3459059.5900N; 322904.4800E); (B) Massive boundstones of late Miocene reef (Koronia Member);  Cross-section
intersecting Polis Basin: (A) SW-NE cross-section intersecting the southern part of Polis Basin from Pegeia region to the Pelathousa
N. Papadimitriou et al.: Oil & Gas Science and Technology - Rev. IFP Energies nouvelles 73, 77 (2018)12
Reef front facies (Fa 2B) evolves into a 20 m-thick unit of
open shelf deposits (Fa 3A; Fig. 6.2A). Open shelf deposits
(Fa 3A) continue up to 280 m (Fig. 8.1A). Biostratigraphy
results have shown that these sediments persist from the
Langhian until the Tortonian and correspond to the
Pakhna Formation. In the Pegeia region, these chalks are
intercalated by reworked material of early Miocene age
(Eaton and Robertson, 1993).
Structural unit 2 (Akamas Peninsula)
Located in the north-western part of Polis Basin, the Aka-
mas Peninsula (Fig. 4) corresponds to a serpentinite belt
which is found on top Mamonia Complex through a series
of thrusts (Figs. 4 and 9.2;Monnet, 2005).
The allochthonous rocks of the Mamonia Complex are
characterized by volcanic breccias (Loutra tis Aphroditis
Formation) sandstones (Vlampouros Formation and
Akamas Sandstones) intercalated with pink calcilutites
and radiolarian mudstones (Episkopi Formation; Fig. 4).
On top of the grey Mesozoic sediments, reef limestones
(Fa 2A) were identiﬁed (Fig. 9). The contact between the
Upper Cretaceous unit and the Miocene reefs is marked
by an erosional surface (Fig. 9). The reefs are 60 m thick
and are mainly composed of boundstone intervals with
corals (i.e.,Porites sp.), and red algae (Fig. 9.1).
To the northern part of Polis Basin (351036.8000N;
3219039.6000E), it has been noticed that the base of the reefs
(0–10 m) consists of red algae fragments, corals, bioturba-
tion and greyish beds of clay. Above this unit, irregular
wavy poritid sheets are alternated with thin marl to pack-
stone beds which could correspond to the Koronia Mb
Structural unit 3 (Polis Graben)
In the ‘‘Polis Graben,’’ Pliocene shallow marine (Payne and
Robertson, 1995), as well as Quaternary continental depos-
its are exposed (Poole and Robertson, 1991,1998,2000).
Hence, to examine the transition from Pliocene to Miocene
units a 260-m deep borehole 1980/076 has been investigated
(Figs. 4,8.2A and 8.4).
The deepest unit in the borehole corresponds open shelf
deposits (Fa 3A) of Miocene age (Fig. 8.2A). The presence
of Cyclicargolithus ﬂoridanus in the chalky limestone indi-
cates a middle Miocene or older age (NN7a or older).
Because some early Miocene age material is reworked, a
younger age (middle Miocene) is preferred.
Fig. 9.  (A) Sedimentary log measured in the northern part of Polis Basin (i.e., Structural unit 2 – Akamas Peninsula;
351036.8000N; 3219039.6000E); (B) Photograph showing the contact between the Koronia Member and the Mesozoic sediments);
 SW-NE cross-section through the northern part of Polis Basin from Akamas Peninsula to the Pelathousa village (see the location
on the map Fig. 4).
N. Papadimitriou et al.: Oil & Gas Science and Technology - Rev. IFP Energies nouvelles 73, 77 (2018) 13
The transition from the Pakhna to Kalavasos formations
(Fig. 8.2A and B) is found at 220 m below ground level
(bgl), where whitish marly chalks pass to the thinly
laminated gypsum (Fig. 8.2B). The Messinian is presented
by 50 m of gypsum that are found between 150–220 m bgl.
Although the description of the borehole is only brief,
outcrop evidence found in the center of the basin, reveal
that the Kalavasos Formation is represented by parallel-
laminated (Marmara) gypsum and massive ﬁne-grained
At 135–150 m bgl, a set of thinly bedded limestones with
allochthonous pebbles marks the end of the MSC. These
are succeeded by 130 m of grey marl with bivalves which
indicate a shallow marine restricted environment
Structural unit 4 (Pelathousa-Peristerona)
The eastern ﬂank of the Polis Basin is referred to as
Pelathousa-Peristerona block (Fig. 4) and is located 10 km
to the east of the Akamas Peninsula (Fig. 4). Geological
observations have shown that Troodos Ophiolites vanish
towards the west (Figs. 4 and 8) whereas late Miocene reefs
are exposed and rest on top of the volcaniclastic sediments of
Kanaviou Formation (Figs. 4 and 8.4).
The Paleogene is not exposed, but ﬁeld observations from
the eastern ﬂank of the basin allowed us to predict that this
unit is composed of chalk alternated with cherts indicating
a deep marine setting (Fa 5B and Fa 5C).
The lower part of the Miocene interval is investigated near
Evretou dam (Figs. 8.3A and 8.4;3458036.3000N,
3228024.6000E). It is composed of alternating highly biotur-
bated bioclastic sand shoal deposits (Fig. 8.3A). The alter-
nation of high energy and low energy facies might
correspond to small cycles that indicate ﬂooding and shal-
lowing of the region. Based on the presence of Cyclicar-
golithus ﬂoridanus (NN7a or older) and the absence of
late Miocene assemblage we propose that these layers repre-
sent a Serravallian sequence.
Moving upwards in the sequence, we observed that the
Miocene unit is composed of massive reefal limestones
(Fa 2A and 2B) which are characterized by thick (3–5 m)
alternated, porous rudstone to grainstones (Figs. 8.3A, C
and 8.4). Based on the results of radiogenic strontium
isotopic analyses, Blanpied (2017) has shown that the
exposed outcrop refers to late Miocene reefs.
4.2.2 Limassol Basin
Located in southern Cyprus, the Limassol Basin is bounded
by the Yerasa Fault System (YFS), the Limassol Forest
Block and the Troodos Ophiolites to the north (Fig. 4).
The Yerasa Fault System is an early to middle Miocene
group of faults (Kinnaird, 2008). The southern boundary
of the Limassol Basin is located offshore Cyprus and is
marked by the Akrotiri High which is now buried under
the Pliocene sediments (Fig. 4). The Akrotiri High is trend-
ing in the same direction with the Yerasa Fault System and
has been described as a thrust system controlled by a
basement high (McCallum et al., 1993). To the west, the
Limassol Basin is separated from Polis Basin by a transverse
zone which extends along the Xeropotamos River (Fig. 4).
Stratigraphic architecture of Limassol Basin
The south-western margin of the Limassol Basin extends
from the Paramali village up to the Petra tou Romiou
(Fig. 4) where white chalks (referred to as Lefkara Forma-
tion) overlie brownish siliciclastic and volcanogenic sedi-
ments (referred to as Mamonia Complex). The basinal
deposits of Paleogene age, in turn, are covered by rudstone
breccias with occasional coral fragments (Fig. 10A).
Although no contact has been found between the Paleogene
and the Miocene formations, we assume that this abrupt
change in the lithology is also marked by a sharp erosional
Based on the composite logs constructed for the south-
ern and the northern part of the Limassol Basin for the
Miocene sequence six stratigraphic sequences have been
recognized, (Figs. 10B and C).Thebaseofeachcycle
corresponds to either an erosional surface or the transition
of very shallow facies (Fa 2, i.e., coral reefs) into slope
deposits (turbiditic deposits; Fa 4). In particular, the ﬁrst
sequence is only present along the southern coast of Cyprus
and is represented by shoal deposits. To the north (Koilani
composite log) and to the east this sequence is marked by
an erosional surface on top of the deep marine sediments
of the Lefkara Formation (Fig. 10C).
The base of the second sequence is found in the
Paramali section is topped by open-shelf deposits
(Fa 3A), whereas in the Koilani section, by peritidal depos-
its. Further up in the Paramali section, this cycle is divided
into three sub-cycles. These sub-cycles represent deepening
and shallowing events, with the alternation of open-shelf
(Fa 3A) and reefal deposits (Fa 2A). Mainly to the south,
the base of sub-cycles 1a, 1b, and 1c are marked at the
top of the coral reef facies and represent the transition from
barrier reef setting to open-shelf and deeper environments.
In the Koilani section, the elementary cycles are absent,
and shoal sands (Fa 1B) are succeeded by open-shelf
(Fa 3A) deposits (Fig. 10C). The Maximum Flooding
Surface (MFS) in both sections are marked by thin, chalky
intervals, which coarsen upwards and end with coral reefs
to the south and bioclastic sands to the north and the east.
To the south, these cycles account for 110 m thickness of
shallow carbonates, whereas to the north and the east there
are only 50 and 60 m, respectively (Fig. 10B).
The initiation of the third sequence in the southern part
of the Limassol Basin is recorded by the transition of
shallow marine facies (Fa 3) into slope and deep marine
sediments (Fa 4 and Fa 5). This change in the depositional
environment is well represented in the south (i.e.,the
Paramali section), where coral reefs (Fa 2A) are overlain
by turbidites (Fa 4A) and pelagic sediments (Fa 5), which
in turn are cut off by MTCs (Fa 4C). The same facies have
been observed to the north and the east (Fig 10B) suggest-
ing an overall drowning. Slope deposits, which are cut off by
N. Papadimitriou et al.: Oil & Gas Science and Technology - Rev. IFP Energies nouvelles 73, 77 (2018)14
MTCs, record the transgressive systems tract, whereas the
MFS is marked by the transition to a deeper setting.
Regarding the thickness, this cycle accounts for 180 m of
sediments to the south, whereas the northern and eastern
parts have 40 and 50 m, respectively, indicating a migration
of the depocenter further to the south. The top of this
sequence is marked by an erosional surface below a remark-
able MTCs (Fig. 10B). Further deepening was recorded in
the fourth sequence which to the north is characterized only
by deformed micritic carbonates.
The following highstand systems tract is represented by
the transition of slope to open-shelf deposits (Fa 3A), which
progressively, evolves, into bioclastic sands. To the north
and east, the slope deposits (i.e., MTCs)passintoopen
shelf (Fa 3A) and relatively shallower sediments. Sea-level
fall and/or uplift can explain the progressive shallowing of
the basin to the south. In this case, the highstand systems
tract and normal regression evolved as a forced regression.
The continuous deposition of shallow marine sediments
observed to the north and east excludes this hypothesis.
The base of the ﬁfth sequence corresponds to another
transgressive surface, above the bioclastic sand to the south,
and below the ﬁrst turbiditic interval in the northern and
eastern parts of the basin. In particular, to the south, this
surface records a further deepening of the area and deposi-
tion of shallow marine sediments. To the north and east,
this cycle is characterized by slope deposits (Fa 4). The
end of this cycle is recorded by the formation of coral reef
Fig. 10. (A) Synthetic cross-section of Limassol Basin. Synthesized measured and interpreted log sections of key exposures of the
Miocene succession in Limassol Basin: (B) Paramali Composite sedimentary log and (C) Koilani composite sedimentary log. They
represent an N-S transect to the western part of the Limassol Basin (see the location of the logs in Fig. 4)
N. Papadimitriou et al.: Oil & Gas Science and Technology - Rev. IFP Energies nouvelles 73, 77 (2018) 15
to the south, and bioclastic sands bars (Fa 1B) to the north
and east (Figs. 10B and C).
The sixth sequence is marked by drowning of the coral
reefs and the bioclastic sands to the south and the north-
east respectively. It seems that the entire basin was ﬁlled
up with shallow-marine/open-shelf deposits. The following
highstand, recorded in the basin by the transition of
open-shelf deposits (Fa 3A) into bioclastic bars with some
patches of coral reef to the south (Fa 2A), and another
bar of bioclastic sands (Fa 1B) to the north and east
(Figs. 10B and C). The bioclastic sands (Fa 1B) to the
south prograde towards the NE whereas those that are
deposited to the eastern part of the basin prograde towards
the NW. The thickness of this cycle varies from the south to
the north. In particular, to the south, these sediments are
180 m thick, whereas the northern and the eastern parts
account for only 60 m of sediments.
5.1 Reconstruction of Polis Basin since the Cretaceous
New offshore studies west of the Polis Basin, based on
seismics and fault plane solution indicate a strike-slip struc-
ture termed the Paphos Transform Fault (Papazachos and
Papaioannou, 1999). These results are in contrast with the
proposed models by Payne and Robertson (1995,2000),
thus prompting a re-evaluation and the proposal of a new
conceptual model that explains the structures and the
deformation encountered during the ﬁeld campaigns in
The early phase of sedimentation in the Polis Basin
occurred in the late Maastrichtian (Ka
¨hler and Stow,
1998) and is marked by pelagic sediments of the Lefkara
Formation. These sediments are found directly on top of
theMesozoic(Fig. 11A). It is assumed that the Polis Basin
was part of a deep basin (Fig. 11A) with the nearest land
being further north and represented by the Taurus
Mountains of southern Turkey (Robertson and Fleet,
1976;Robertson et al., 1991,2012).
From the late Oligocene to the early Miocene, the
Eastern Mediterranean experienced a regional uplift coeval
to a long-lived lowstand (Haq et al., 1988) which might be
also linked to the collision of the African and Eurasian plates
(Dercourt et al., 1986;Robertson, 1998a, b;Jolivet and
Faccenna, 2000;Dargahi et al., 2010). The inferred uplift
in the Polis area is identiﬁed by blind thrusts and karstiﬁca-
tion on top of the Lefkara Formation that is exposed locally,
near the Kathikas village (Fig. 11B). It seems that the Upper
Cretaceous thrust activity at the Troodos Ophiolites
propagated further to the south-west near Kathikas village
and thus a new basin was formed (Fig. 11B).
During its early stages, the Polis Basin was character-
ized by reefal (Fa 2A) and reef front (Fa 2B) sediments that
evolved into open shelf deposits (Fa 3A; Fig. 11B). Cross-
sections in the basin portray the accurate pattern of the
early Miocene reefs which in turn reﬂects the topography
of western Cyprus during that time (Fig. 4). These reefs
(i.e., Terra Member) were mainly formed on the basin
margins. Reworked material from Cretaceous (Kathikas
Fm) and Oligocene (Lefkara Fm) formations are found in
the early Miocene sediments depicting a tectonically active
period. The reef front facies were identiﬁed in an NW-SE
trending zone parallel to the linear trend of the reef patches
(Fig. 11B). Although there are no signs for the direction of
the paleoslope, the borehole data suggest that the high-
energy reef front sediments evolved into low energy shelf
deposits in the center of the Polis Graben. During the
Langhian, a global sea-level rise occurred (Haq et al.,
1988;Hawie et al., 2013). In the Polis Basin, the bioclastic
shallow marine sediments with reef front and reefal facies
were succeeded by open shelf deposits (Fig. 11C).
Microscale analysis revealed that the middle Miocene
sediments were ﬂoored with reworked early Miocene mate-
rial (i.e., C. fenestratus,Z. bijugatus,andS. dissimilis).
Sea-level rise alone cannot explain the deepening of this
region, and it is unlikely that no tectonic activity was
involved in the sedimentary evolution of the basin during
this period. Payne and Robertson (2000), suggest that the
dominant deformation mechanism was associated with an
extensional regime expressed by large normal faults which
bound of the basin ﬂanks. This period of extension may
be associated with slab rollback along the northward
subduction of the African plate during the Miocene. Indeed,
during our ﬁeld studies, we observed east of the Kathikas
high/plateau NNW-SSE normal faults cutting Pakhna
sediments (Fig. 4). These faults have been interpreted by
Monnet (2005) as pronounced escarpments associated with
gravitational forces and not as deep structures.
A possible mechanism for the subsidence of the basin
can be the tilting of its basement which might have
occurred due to the SW propagation of the thrust belt
and the contemporaneous uplift of the southern slopes of
the Troodos Mountains. As a consequence, some gravita-
tional faults were created in the center of the basin which
was covered by relatively deeper sediments. In contrast,
the north-eastern ﬂank of the basin (near the slopes of
Troodos Mountains) was covered by sand shoal deposits
which were prograding to the south-west (Fig. 11C).
Nevertheless, the local accommodation space in
piggyback basins can be created by out of phase thrust
activity and the subsequent migration of the depocen-
ter (Zoetemeijer et al., 1993;Chanvry et al., 2018). Usually,
the depocenters in piggyback basins migrate towards to
the foredeep as a result of the propagation of the thrust
belt. Out of phase thrusting (back thrust) can lead to
the migration of the depocenter towards the inner land.
Thus, further investigation should be undertaken in order
to examine the mechanisms of subsidence during this
During the late Tortonian, reefs (Koronia Member)
colonized the eastern and the western ﬂanks of the Basin
(Fig. 11D). The development of the Koronia Member
should be connected with a late Miocene thrust activity
which propagated towards the SW (Pafos Thrust) and/or
the sea-level drop before the MSC (Haq et al., 1988). This
thrust activity is thought to be ceased during this period
since the thrust faults in the Pelathousa Region are sealed
by the Koronia Member (Fig. 11D).
N. Papadimitriou et al.: Oil & Gas Science and Technology - Rev. IFP Energies nouvelles 73, 77 (2018)16
In the eastern ﬂank, no extensional traces prior to Pleis-
tocene have been identiﬁed. Nevertheless, the presence of
Messinian evaporites on top of the neritic carbonates of
Pakhna Formation as well as the thick sequence (150 m)
of Pliocene marine sediments in the center of the basin
(Payne and Robertson, 1995), led us to propose the forma-
tion of a N-S asymmetrical depression during the post
Tortonianperiod(Fig. 8.4). This NW-SE depression formed
Fig. 11. Synthesized scheme and paleoenvironmental maps illustrating the evolution of the Polis Basin from the Late Cretaceous
until the Pliocene.
N. Papadimitriou et al.: Oil & Gas Science and Technology - Rev. IFP Energies nouvelles 73, 77 (2018) 17
within the present-day Polis valley (Structural unit 3;
Fig. 8.4;Follows, 1992;Payne and Robertson, 1995)and
was controlled by normal gravitational faults that created
during the tilting of the basement due to the propagation
of the thrust belt (Fig. 11E).
Field observations have shown that the Kalavasos
Formation (Messinian evaporites) is absent from the north-
ern part of the Polis Graben (Kinnaird, 2008). In addition,
at the eastern and the western ﬂanks of the graben, the
Miocene carbonates are directly overlain by the Pleistocene
deposits without any evidence of Messinian or Pliocene
sedimentation (Figs. 4 and 11E). In contrast, the southern
part of the graben in the center of the basin was deep
enough to be ﬁlled with the Messinian evaporites as indi-
cated by the presence of outcropping laminated gypsiferous
marls and gypsum in the boreholes (Fig. 8.2).
During the Pliocene, the deposition of evaporites was
followed by the inundation of the basin with sea waters
(Robertson, 1998a;Bowman, 2011;Hawie et al., 2013;
Gorini et al., 2015). The increased thickness of the Nicosia
Formation towards the center of the Polis Graben (Figs. 8.4
and 9.1) and its absence from the Pegeia region and
Akamas Peninsula reﬂects the inﬁll of the existing basin
The deposition of shallow water calcarenites within the
Polis Graben during the late Pliocene and early Pleistocene
combined with the non-deposition elsewhere reﬂects the
ﬁrst stages of the general uplift of Cyprus (Payne and
Robertson, 1995;Kinnaird, 2008). The continuous shallow-
ing during this time can be linked with the southward
propagation of the thrust activity and thus the migration
of the depocenter towards the foredeep. It is assumed that
the youngest thrust-sheets are located offshore and are
the equivalent of the Cyprus Arc.
5.2 Reconstruction of Limassol Basin
Based on biostratigraphic results and ﬁeld observations, it is
suggested that the Limassol Basin was created during the
early to middle Miocene (Burdigalian – Langhian). During
this time, compression and propagation of the thrust belt
to the south, resulted in the initiation of several thrusts, in
an NW-SE direction (i.e., the Yerasa and Akrotiri fault
systems), bounding the Limassol Basin (Fig. 4). The sedi-
mentary record shows an unconformity to the north of the
basin, where basinal deposits (contourites) are directly over-
lain by intertidal sediments (Figs. 12A and B). This abrupt
shallowing of the region marks an episode of uplift that on a
regional scale, might be attributed to the position of Cyprus,
in the complex zone of convergence between Africa and
Eurasia (Robertson and Woodcock, 1986;Eaton and
Robertson 1993). To the south, the basin was colonized by
small, mounded bioherms, surrounded by reef-talus and
shallow-water carbonates whereas to the north by bioclastic
shoals (Figs. 12B and C;Stow et al., 1995). It is assumed
that the reefs in the southern margin of the basin colonized
the relief topography that was created by a blind thrust
(referred to as Akrotiri High or Akrotiri Fault System;
During the Tortonian, the Limassol Basin experienced
signiﬁcant subsidence and the Miocene reefs to the south
as well as the bioclastic sands to the north were covered
by slope to deep marine sediments (Fig. 12D). Progressive
thickening of the Tortonian sediments, from north to south,
suggests a southward migration of the depocenter
(Fig. 12D). In particular, to the south, 500 m of slope to
basinal deposits have been measured, whereas the northern
part of the basin accounts for only 80 m of sediments. The
subsidence of the basin can be explained by the thrust
activity at the southern slopes of the Troodos Mountains.
The early Miocene Yerasa Thrust System was still active
during the Tortonian and as a basement thrust caused the
displacement of the Troodos Ophiolites to the south and
thus the thickening/stacking of the crust. Consequently,
the northern part of the Limassol Basin was out of isostatic
equilibrium and must have partially sunk. Therefore, the
southern part of the basin which was underlined by rela-
tively thin crust compared to the north, subsided, allowing
further subsidence of the overlain sedimentary basin.
On the ﬂanks of the basin, MTCs have been identiﬁed.
Previous works suggest that the MTCs are oriented in east-
southeast (Farrell and Eaton, 1987;Eaton and Robertson,
1993;Kinnaird, 2008), or in the east-west direction (Lord
et al., 2009). During the present study, measurements along
the Limassol-Pafos highway to the south (3441012.5100N,
3241024.6400E), and near Dora village to the north
(3445014.4200N, 3241019.1200E), agree with Eaton’s (1987)
results (Fig. 4). However, it is unlikely that the slope was
dipping uniformly away from the Troodos Ophiolites, since
Eaton and Robertson, (1993) noted two exceptions near
Agia Fila and Happy Valley, depicting a palaeoslope
direction towards the north (Fig. 4). Thus the direction of
the slope must have also been controlled by a northward
dipping structure probably located south of Cyprus. This
structure could correspond to the Pafos Thrust (Monnet,
2005). Since the MTCs are the result of intense seismic
activity (e.g., earthquakes) (Alves, 2015;Arfai et al.,
2016;Guan et al., 2016) it is proposed that, during the
Tortonian, the thrust belt propagated further to the south
resulting in the initiation of the Pafos Thrust (Fig. 12E).
Before the Messinian, the Limassol Basin was ﬁlled with
shallow marine sediments (Fig. 12E) and therefore the
basin experienced an uplift probably associated with the
southward propagation of the thrust belt. Along the eastern
margin of the basin, bioclastic sand shoals began to pro-
grade towards the NW, whereas to the south the bioclastic
sands were prograding towards the NE (Figs. 10B and C).
5.3 A new model for the Cenozoic tectonostratigraphic
evolution of Cyprus
Using the ﬁeld data presented in this contribution, and
recently published geophysical data (Welford et al., 2015;
Reiche et al., 2016;Granot, 2016) a new paradigm for the
tectonostratigraphic evolution of southern Cyprus since
the Late Cretaceous can be proposed.
A northward-dipping subduction zone is thought to
have been initiated during the Turonian, with the opposing
movement of the Afro-Arabian and Eurasian plates, result-
ing in the formation and obduction of the Troodos
Ophiolites (Bowman, 2011;Montadert et al., 2014). After
their obduction (Campanian until the Maastrichtian), the
N. Papadimitriou et al.: Oil & Gas Science and Technology - Rev. IFP Energies nouvelles 73, 77 (2018)18
ophiolites juxtaposed with the Mamonia Complex
(Fig. 13A;Bailey et al., 2000). At that time, westerly
directed thrusting, at the slope of Troodos Mountains,
and sinistral transpression along the Mamonia Suture
Zone took place (Figs. 1 and 2). The transpressional move-
ment is recognized by a transverse fault zone along the
Xeropotamos River (Fig. 4). This fault zone deﬁnes the
limit of the Mamonia Complex onshore Cyprus and has
been interpreted as the prolongation of the Continental
Oceanic Boundary (COB), that is identiﬁed offshore
Cyprus (Figs. 1,2and 4;Granot, 2016).
From the Maastrichtian until the late Eocene, the
Mediterranean was under deep marine conditions, with no
evidence of ongoing plate convergence (Fig. 13B;Robertson
et al., 2012). During the late Eocene, the convergent plate
boundary migrated southwards, following the collision of
the Keryneia Range with the Troodos Ophiolites
(Fig. 13B;Robertson and Woodcock, 1986). The collision
and the southward propagation of the thrust activity
resulted in the submergence of the Keryneia Range and
the subsequent development of an extensive submarine
fan system (Robertson and Woodcock, 1986).
Fig. 12. Synthesized scheme and paleoenvironmental maps illustrating the evolution of the Limassol Basin during the Miocene.
N. Papadimitriou et al.: Oil & Gas Science and Technology - Rev. IFP Energies nouvelles 73, 77 (2018) 19
Fig. 13. Paleoreconstructions showing the evolution of southern Cyprus based on the forward propagation model.
N. Papadimitriou et al.: Oil & Gas Science and Technology - Rev. IFP Energies nouvelles 73, 77 (2018)20
In southern Cyprus, the sedimentation style shows a
diachroneity between the Limassol and Polis basins,
suggesting a differential uplift. For instance, the southern
margin of Polis Basin was controlled by thrust faults, where
small reefs formed (Terra Member), suggesting southward
propagation of the thrusts from the southern slopes of the
Troodos towards the Kathikas village (Fig. 13B). On the
other hand, in the Limassol Basin, a relatively deeper
marine environment prevailed until the early Miocene
(Fig. 13B). During the following period (Langhian –
Serravallian), the Limassol Basin was under a shallow
marine setting (Fig. 13C). Eaton and Robertson (1993)
suggested that the uplifted area was controlled by south-
ward verging thrusts (YFS) that are visible on the slopes
of the Troodos Mountains.
Evidence of Late Miocene compressional deformation is
documented in both basins. In the western ﬂank of Polis
Basin, an erosional surface between Late Miocene reefs
and the Cretaceous pillow lavas depicts an early/Late
Miocene movement along a thrust fault. In the eastern ﬂank
of the Polis Basin, the Pelathousa-Peristerona thrust was
active as evidenced by the deposition of the Koronia
Member reefs which are tilted towards the east and are in
direct contact with the Campanian sediments (Fig. 13D).
On the contrary, the Limassol Basin suffered by signiﬁcant
subsidence and deep marine sediments deposited in the
southern part of the basin (Fig. 13D).
The collision between Eratosthenes Seamount and
Cyprus Arc during the latest stages of the Miocene
(Robertson, 1998b;Papadimitriou et al., 2018) resulted in
a progressive shallowing of the two Neogene basins
(Fig 13E). During this time, pre-existing normal faults con-
tinued to control the central part of the Polis Basin, (Polis
Graben) whereas, on its ﬂanks, small patch reefs were devel-
oped (Fig. 11). In the Limassol Basin, the sedimentation was
dominated by prograding bioclastic sands with some patch
reefs in the eastern and the southern margins (Fig. 12D).
New Geophysical data that cover the northern part of
the Levant Basin support that during the late Miocene
compressional nature of the Latakia Ridge system changed
to a sinistral strike-slip (Hall et al., 2005a,b). This change
in the compressional regime between the two plates, may
have resulted in the westward propagation of the COB
(Fig. 1) explaining the eastward migration of the apparent
deformation observed onshore Cyprus (Fig. 13E).
The Plio-Pleistocene time is envisaged as a period of
tectonic uplift (Sage and Letouzey, 1990;Orszag-Sperber
et al., 1989;Eaton and Robertson, 1993;Robertson,
1998b;Kinnaird and Robertson, 2012). Evidence of this
uplift has been found within the Polis and Limassol basins
(Fig. 13F). In particular, late Pliocene and early Pleistocene
shallow water calcarenites were observed in the center of
the Polis Basin (i.e., Polis Graben) whereas its marginal
areas were subaerially exposed (Fig. 13F;Payne and
Robertson, 1995;Kinnaird, 2008).
5.4 Piggyback vs. ﬂexural basins
Both Limassol and Polis basins were formed under a com-
pressive regime (Figs. 11–14). The transition from one basin
to the other was previously deﬁned by a slightly oblique
ramp outcropping from the coastline to the southern slopes
of the Troodos Mountains, through the Xeropotamos River
(Fig. 4) trending in a north-south direction (Monnet, 2005).
¨hler and Stow, (1998) suggested that gradual uplift
and an increase in sediment input from the northeast can
explain this lateral change in depositional environment
between the two basins. The question that arises from this
statement is: What might be the cause of the progressive
and differential uplift between the two basins?
A good explanation would be given due to the
structured part of the Polis Basin substratum associated
with the collision of the Mamonia Complex with the
Troodos Ophiolites (Bailey et al., 2000;Lapierre et al.,
2007). However, this hypothesis does not explain the
diachroneity and the spatial distribution of Miocene thrusts
alongside the Limassol and Polis basins (Figs. 14 and 15).
In the Polis Basin, the early Miocene thrust sheet
extends from Mesogi to the Akamas Peninsula, whereas
in the Limassol Basin, this thrust is referred to as Yerasa
Fault System (Figs. 14 and 15). It is proposed that the
two faults are connected through the transverse fault zone,
across the Xeropotamos River (Fig. 4).
The Polis Basin is formed on top of the sedimentary
cover of the Mamonia Complex (Fig. 14). In contrast, the
Cenozoic sediments found in the Limassol Basin overly
the rigid ophiolitic complex (Fig. 14).Thenatureofthe
substratum and the ‘‘subducting’’ crust between the two
basins (oceanic in Polis and thin continental in Limassol)
can explain the diachroneity of the thrusts as well as the
differences in intensity of the deformation style between
the two. As one of a series of piggyback basins, the Polis
system must have been formed in association with upper
crust, thin-skinned NW-SE thrusts, detached from their
basement along a decollement horizon (Zoetemeijer et al.,
˜oz et al., 2013). It is expected that the decolle-
ment level of such thrusts is within the weak sediments of
the Mamonia Complex (Fig. 14). In contrast, 500 m of
Tortonian series in the Limassol Basin implies the local
emplacement of a substantial lithospheric load to generate
such rates of ﬂexural subsidence.
Hence, thin-skinned vs. thick-skinned tectonics can
explain the eastward increase (from 50 to 500 m) in the
sedimentation, along the southern limits of the Polis and
Limassol basins. The southward propagation of the thrust
belt in both basins and the synchronization of the deforma-
tion is recorded by the Pafos Thrust (Fig. 15A). In the Polis
Basin, the activity of this fault resulted in the uplift of the
ﬂanks and the subsequent tilting of the basement and thus
the formation of Polis Graben. In contrast, in Limassol
Basin, the initiation of Pafos Thrust is recorded by MTCs
which cut the deep marine sediments. It has been
postulated that the MTCs record also the collision of the
Eratosthenes Carbonate platform with the Cyprus Arc
(Papadimitriou et al., 2018).
A good analog of the Polis and Limassol basins would be
the Graus-Tremp-Ainsa basins which are located in the
´es (Chanvry et al., 2018). In particular, in
the south Pyrenees, the easternmost basin (i.e., the
Organya-Tremp-Ager sub-basins) is controlled by three
N. Papadimitriou et al.: Oil & Gas Science and Technology - Rev. IFP Energies nouvelles 73, 77 (2018) 21
Fig. 14. Synthesized chronostratigraphic chart which shows the lithologies and the major hiatuses in Limassol and Polis basins with
respect to the main geodynamics.
N. Papadimitriou et al.: Oil & Gas Science and Technology - Rev. IFP Energies nouvelles 73, 77 (2018)22
Fig. 15. [A] Geological map of southern Cyprus with the main structures that have been discussed during the present study;
(A) Cross-section of Polis Basin; (B) Cross-section of Limassol Basin; CA – Cyprus Arc; PFS – Pafos Fault System; ATF – Arakapas
Transform fault; PB – Polis Basin; LB – Limassol Basin; KF – Kathikas Thrust; [B] Sketch showing the main structural units of the
central Pyrenees and the major anticlines of the Ainsa Oblique Zone at the western boundary of the major thrust in the central
Pyrenees. Crustal cross-sections at both sides of the Ainsa Oblique Zone illustrate changes of the structural style along strike.
AOZ – Ainsa Oblique Zone; A – An
˜isclo Anticline; B – Boltan
˜a Anticline; M – Mediano Anticline; C – Cotiella; PM – Pen
˜esa; SCU – South Pyrenean Central Unit; Bx – Bo
´ixols; SM – Serres Marginals (modiﬁed from Mun
˜oz et al., 2013).
N. Papadimitriou et al.: Oil & Gas Science and Technology - Rev. IFP Energies nouvelles 73, 77 (2018) 23
thrusts creating a series of piggyback basins (the Bo
Montsec and Sierra Marginales) whereas to the west (Ainsa
to Jaca basins), a ﬂexural basin prevails; bounded by one
thrust-front zone to the south (Fig. 15B;Chanvry et al.,
˜oz et al., 2013). The two systems are separated
by a lateral ramp (the Mediano structure, Fig. 15B)and
have been set up during the initiation of the convergence
between Iberia and Eurasian plates (Chanvry et al., 2018;
Fillon et al., 2013;Mun
˜oz et al., 2013). The partitioning
of the basins in south Pyrenees can be explained by the
different nature of the substratum below the Paleogene
sediments in the two areas (Fig. 15B).
As it is shown in Fig. 15B in the Axial Zone, the
Mesozoic and Cenozoic sediments are ﬂoored by Triassic
evaporites whereas to the west these evaporites are absent.
South of the Axial Zone, the Mesozoic cover is detached
along the Triassic evaporites forming the Bo
and Sierra Marginales thrust-sheets. In the western basins,
the Mesozoic and Paleogene covers are coupled to the base-
ment, and the basement thrusts are mostly imbricated
(Fillon et al., 2013;Mun
˜oz et al., 2013). Therefore the
central zone was controlled by thin-skinned tectonics and
the western part by thick-skinned. The thin-skinned system
propagated and migrated both southward and westward in
response to the non-orthogonal collision of Iberia with
Europe during Paleogene mountain building. In contrast,
to the west, the thrust system has been interpreted to be
a longer-lived structure that initiated during the exten-
sional tectonic regime in Middle Cretaceous time and
inverted during the main episode of the Pyrenean collision
(Fillon et al., 2013;Mun
˜oz et al., 2013). The development
of the basement thrust system caused regional subsidence
along the south Pyrenean foreland margin which was
subsequently halted by local uplift associated with the
west-migrating thin-skinned thrust system (Chanvry
et al., 2018;Fillon et al., 2013;Mun
˜oz et al., 2013).
Based on ﬁeldwork in southern Cyprus, and the review of
previous works, a better understanding of the tectonostrati-
graphic evolution of Cyprus has been achieved. The results
pertaining to the sedimentary ﬁlling and the structural
evolution of two Neogene basins (Polis and Limassol) as
dThe Polis Basin is interpreted as a piggyback basin
that is controlled by thin-skinned tectonics. This kine-
matic evolution of the thrust belt is localized by the
weak sediments of the Mamonia Complex below the
detached Cenozoic successions. The Polis piggyback
basin has evolved from the interaction of four thrusts:
(a) the Troodos thrust, (b) the Kathikas thrust,
(c) the Pafos Thrust and (d) the Cyprus Arc.
dThe geometry and thus the sedimentation in Limassol
Basin was controlled by a main blind-front thrust to
the south and an oblique ramp to the west. The
forward propagation of the thrust belt is only related
to crustal shortening which recorded the main
geodynamics. The main thrusts that control the
Limassol Basin are (a) the Yerasa Fault, (b) the Pafos
Thrust and (c) the Cyprus Arc.
dThe slightly oblique ramp (referred as a transversal
zone along the Xeropotamos River) determines the
transition between the shallow marine sediments
deposited in Polis Basin and the slope to basinal
setting observed in the Limassol Basin.
dThe MTCs identiﬁed in the Limassol Basin are
associated with an intense tectonic activity that
occurred during the late Miocene. Similar MTCs,
but with different scale have been identiﬁed to the
eastern and the western sites of Eratosthenes
Seamount offshore Cyprus.
dThe integration of structural and sedimentological
data to propose a reconstitution of the northern
margin of the Cyprus Arc between Oligocene to late
Miocene. The kinematic model that best matches all
the observed data is that of the forward propagation
Acknowledgments. The authors would like to acknowledge the
Ministry of Energy, Commerce, Industry and Tourism of the
Republic of Cyprus as well as the Geological Survey Department
of Cyprus. A very warm thank you to Dr. Zomenia Zomeni and
Dr. Efthymios Tsolakis for their contribution and their help dur-
ing the ﬁeld studies onshore Cyprus.
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