Some 20 Myr after the Late Jurassic to Early Cretaceous
obduction and collision at the eastern margin of Adria, the eroded Pelagonia
(Adria)–Axios/Vardar (oceanic complex) contact collapsed, forming the
Kallipetra Basin, described around the Aliakmon River near Veroia (northern
Greece). Clastic and carbonate marine sediments deposited from the early
Cenomanian to the end of the Turonian, with abundant olistoliths and slope failures at
the base due to active normal faults. The middle part of the series is
characterized by red and green pelagic limestones, with a minimal contribution
of terrigenous debris. Rudist mounds in the upper part of the basin started
forming on the southwestern slope, and their growth competed with a
flux of ophiolitic debris, documenting the new fault scarps affecting the
Vardar oceanic complex (VOC). Eventually, the basin was closed by
overthrusting of the VOC towards the northeast and was buried and heated up
to
The Hellenides are an integral segment of the main Alpine–Himalayan orogenic
belt (Fig. 1). They have recorded polyphase Alpine deformation since the
Middle Jurassic, when they were involved in the obduction of imbricate
oceanic units over the eastern Apulian margin (e.g., Bernoulli and Laubscher, 1972; Zimmerman and Ross, 1976; Schmid et al.,
2020). In the Internal Hellenides (Fig. 1), continuous convergence led to
the collision of continental promontories with Eurasia in the Late
Jurassic–Early Cretaceous that built a metamorphic crustal-scale orogenic
wedge involving the Pelagonian zone and Rhodope (Ricou et al., 1986; Burg et al., 1996; Schenker et al., 2014). In the Late Cretaceous, the metamorphic
thrust sheets of the Pelagonian zone were exhumed to shallow depths. This is
testified by a cooling below ca. 240
Location of the Hellenides and study area in the Alpine Mediterranean chain. Modified from Burg (2012).
In the Pelagonian zone and adjacent units, the record of this orogenic system in the time interval between collision in the Early Cretaceous and resumed thrusting in the Late Cretaceous–Early Cenozoic remains sparse, and the discontinuous, sometimes contrasting, large-scale interpretations stem from the difficulties in establishing a coherent tectonic history across the Rhodope and the Pelagonian zone (Fig. 2). To elucidate part of these controversies, this study investigates the small Upper Cretaceous Kallipetra Basin that formed on both oceanic and continental units along the eastern Pelagonian margin (Fig. 1) and was overthrusted by serpentines of a Jurassic oceanic floor (the ophiolitic fragments lying on the Pelagonian zone named Axios/Vardar/Almopias zone by Schenker et al., 2015). According to the scenarios proposed in Fig. 2, the Kallipetra Basin may have formed (i) within a long-lived Jurassic–Cretaceous passive margin characterized by the income of flysch from the approaching Rhodopian trench to the north–northeast (e.g., Papanikolaou, 1989; Ricou et al., 1998; Papanikolaou, 2009); (ii) during an extensional tectonic event in between the Pindos obducting from the south–southwest and the Axios/Vardar/Almopias zone subducting to the north–northeast (Sharp and Robertson, 2006); (iii) over the obducted Axios/Vardar/Almopias zone (Froitzheim et al., 2014); or (iv) within a collisional wedge that incorporated the obducted Axios/Vardar/Almopias zone (Schenker et al., 2014).
The Kallipetra Basin collected coarse detritus – metamorphic and ophiolitic
rocks from its shoulders – followed by the deformation of its deposits, by
thermal conditions that locally partially or totally reset the cooling ages,
and by cooling during the Late Cretaceous (Schenker et
al., 2015). However, the stratigraphic evolution and the depositional age of
this basin are so far only partially constrained. Moreover, it remains
unclear how thermal conditions (temperatures
This study uses conventional geological mapping techniques, stratigraphic analysis, illite crystallinity, and low-temperature thermochronology to obtain new constraints on the tectonic evolution of the eastern margin of the Pelagonian zone and to unravel the Late Cretaceous detrital record. Our data indicate that the Upper Cretaceous Kallipetra Basin was relatively shallow and tectonically active as testified by the presence of olistoliths, large gravitational features such as rotational growth faults and slumping, and early diagenetic deformation. Rudist bioherms were accumulated on the slopes of the basin with flank deposits dipping into the basin. The bioherms were terminated through environmental restriction or burial due to increased serpentinite sediment input from the south–southwest-eroding oceanic complex. Moreover, based on illite and petrographic data, we find an inverted, high, non-linear geothermal gradient related to a heating event, which likely occurred during the overriding of the Vardar oceanic complex (VOC) in the Late Cretaceous.
Following the Variscan Orogeny and Permian strike-slip and extension (Schenker et al., 2018), the Permian–Triassic rifting led to the creation of the Tethys and its seaways, namely the Pindos, Vardar/Maliac, and Meliata basins, which continued to open during the Triassic to Early Jurassic (e.g., Bernoulli and Laubscher, 1972; Schmid et al., 2008; Papanikolaou, 2009). The convergent motion between Eurasia and Adria led to a northward intra-oceanic subduction in the Vardar in the Early–Middle Jurassic that saw the production of magmatic arcs to the north (Dimitrijevic, 1982; Bortolotti et al., 1996; Burg, 2012). In the Late Jurassic, there was southwestward obduction of the Tethys ophiolite from the Vardar Ocean onto the passive continental margin of the Pelagonian zone to the south (Bernoulli and Laubscher, 1972; Dimo-Lahitte et al., 2001). Jurassic-to-Lower-Cretaceous sediments were imbricated during the accretion of the ophiolitic units (e.g., Robertson and Dixon, 1984; Bortolotti et al., 2005; the complex named Axios/Vardar/Almopias zone in Schenker et al., 2015). Continuous crustal shortening caused the accretion of Rhodope by the latest Jurassic–Early Cretaceous and of the Pelagonian zone by the Early Cretaceous (Figs. 1 and 2; Burg et al., 1996; Ricou et al., 1998; Schenker et al., 2014; Moulas et al., 2017). The buried Pelagonian basement experienced regional amphibolitic-facies metamorphism to the north (U–Pb zircon ages from the leucosomes of migmatites at 130–117 Ma; Schenker et al., 2015, 2018) and an upper greenschist- to blueschist-facies metamorphism to the south (Ar–Ar ages on muscovite at 100–85 Ma; Schermer et al., 1990; Lips et al., 1998).
The tectono-sedimentary history during the Early Cretaceous is highly
debated. Sharp and Robertson (2006) suggest that the Pelagonian zone and its
emplaced ophiolitic rocks underwent extensional exhumation already during
the Late Jurassic. Rather, Early Cretaceous tectonic activity has been
recorded near Edhessa (Fig. 1) in the Axios/Vardar/Almopias zone, along with
late Aptian–Early Albian transgression on both the Pelagonian platform and
the Axios/Vardar/Almopias zone (Mercier, 1968; Mercier and Vergely, 2002).
To the south, the Aptian–Albian time was characterized by a sedimentary
hiatus (
Thereafter, transgressive Cenomanian-to-lower Campanian limestones and deep-water Paleocene turbidites unconformably overlay the eroded Pelagonian and Axios/Vardar/Almopias imbricated units (Mercier, 1968; Mercier and Vergely, 2002; Papanikolaou, 2009) attesting deepening below sea level of the Rhodope–Pelagonian crustal-scale orogenic wedge. Moreover, during the Late Cretaceous to Eocene and locally since the Campanian, the imbrication of the Axios/Vardar/Almopias units resumed at several locations in relation to thrusting with vergence to the NE and to the SW. This has been documented in the central–eastern Vardar near the study area (Paikon Window; Godfriaux and Ricou, 1991; Bonneau et al., 1994; Brown and Robertson, 2003; Katrivanos et al., 2013), in the northwestern Vardar (Grubić et al., 2009; Ustaszewski et al., 2009), in the northeastern Pelagonian zone (Kilias et al., 2010), in the southern Pelagonian zone (Baumgartner, 1985), and in the Pindos zone (e.g., Aubouin, 1959, 1973; Papanikolaou, 1997). Continuous convergence up to the Neogene progressively deformed the continental margin of the Adriatic plate into southwest-verging fold and thrust sheets (Fig. 1; Channell and Hovarth, 1976). Final exhumation of the stacked crustal and oceanic slices occurred through extensional metamorphic domes between the Eocene in the north and late Neogene in the south (e.g., Lister et al., 1984; Dinter and Royden, 1993; Gautier et al., 1993, 1999; Brun and Sokoutis, 2007; Jolivet and Brun, 2010; Burg, 2012).
The Pelagonian basement consists of deformed (i) orthogneisses crosscut by
leucogneisses and leucogranites; (ii) mafic amphibolite bodies; and (iii) interlayered marbles (Schenker, 2013, and references therein).
Cooling of the Pelagonian core complex carapace rocks may have started at or
after collisional doming at
The Axios/Vardar/Almopias unit includes a mélange zone made of tectonically superimposed marbles, serpentinites (ophicalcites), flysch–phyllitic series, volcanoclastic sediments, amphibolites, and carbonatic sequences imbricated southwestward during the Late Jurassic obduction over the Pelagonian zone (e.g., Smith et al., 1975; Ricou and Godfriaux, 1995; Sharp and Robertson, 2006; Ferriere et al., 2016). In the study area, the Axios/Vardar/Almopias unit is represented by serpentinites that are referred to as the VOC, which consists of at least five lithologies: (i) ophicalcites; (ii) dark massive fractured to brecciated serpentinites; (iii) sedimentary serpentinite breccias; (iv) sedimentary serpentinite breccia with platform carbonate clasts; and (v) foliated serpentinites and limestones. Ferromanganese-rich chert nodules within the VOC, dated further to the south at approximately 175 Ma by Chiari et al. (2013), attest the involvement of this Jurassic oceanic floor in the intra-oceanic Tethys subduction and subsequent obduction.
On the eastern margin of the Pelagonian zone, relatively thick packages of
Upper Cretaceous carbonate and siliciclastic sediments with both a
Pelagonian and ophiolitic provenance unconformably cover the VOC and the
Pelagonian basement (Sharp
and Robertson, 2006; Papanikolaou, 2009; Schenker, 2013; Schenker et al.,
2015). In the study area, the sediments belong to a sedimentary basin that
here is referred to as the Kallipetra Basin (Fig. 3). It formed as an
elongated NNW–SSE-oriented belt overlying the VOC and the Pelagonian
continent. In this basin, the presence of reworked Lower Cretaceous
orbitolinids,
Geodynamic interpretations of the Hellenides in the Early Cretaceous according to different authors. Note that there is no consensus on the Early Cretaceous geodynamic framework at the onset of the Kallipetra Basin between the Pelagonian zone and the Vardar domain.
Geological map of the study area. Yellow circles indicate locations of illite samples; red circles indicate location of ZFT samples and their respective ages.
From the early Oligocene, an overall southwestward tectonic denudation from shallow depths is documented in the Kallipetra Basin by AFT ages of 32.7 Ma (sample 10-128) and in the Pelagonian basement by ZFT ages of 24–20.7 Ma and AFT ages between 22.9 and 16.1 Ma (Schenker et al., 2015).
Geological mapping and structural analysis were conducted to reconstruct the
geometry of the basin and the ductile and brittle deformation that affected
the Kallipetra Basin and the overlying VOC. The paleogeography, depositional
environment, and age of the sedimentary sequences were determined based on
stratigraphic logging, optical microscopy, and biostratigraphy. Planktonic
foraminifera and nannoplankton were used to establish ages of the
sedimentary succession. Simple smear slides were produced using standard
techniques to retain the nannofossil assemblages and original sediment
composition. Quantitative analyses were carried out using a polarizing light
microscope at a magnification of
The Kübler index of illite crystallinity is a method used to determine
the diagenetic grade in metapelitic sequences by measuring the changes in shape
of the first dioctahedral illite–muscovite basal reflection at a 10 Å
X-ray diffraction (XRD) spacing (Kübler and
Jaboyedoff, 2000). To analyze illite crystallinity, bulk-rock mineralogy was
obtained through the conventional powder XRD method using the ARL Thermo
X'tra powder diffractometer at the University of Lausanne. Samples were then
de-carbonated, followed by the extraction of
Significant asymmetrical peak broadening, caused by a tail in the 10 Å
peak and produced by the presence of smectite and expandable mixed layers,
is reduced following EG treatment (Abad, 2007). These peaks may
indicate the presence of detrital illite, which gradually decreases with
burial and essentially disappears in the anchizone
(Kübler and Jaboyedoff, 2000). The decrease in
KI values with increasing metamorphic conditions and temperatures is a
consequence of the increase in the number of layers and the disappearance of
expanding layers. The Neuchâtel IC scale was calibrated with the
Lausanne diffractometer and therefore produced anchizone limits of
0.18 and 0.36
Two of five collected samples provided enough zircons to date using ZFT analysis: V1503 and V1504. These samples integrate our
previous samples (10-128, 10-129, 10-130; Schenker et al., 2015). The new
samples were collected with the aim of revealing the full age distribution,
which in our previous samples was limited by the low number of available
zircons. To this goal, the new samples were
The study area is divided into three units: (1) the Pelagonian basement; (2) a stratigraphic unit that we name the Kallipetra Formation, described here for the first time; and (3) the VOC. The Kallipetra Formation is the focus of this study and consists of several lithofacies that collectively characterize a sedimentary basin (Fig. 3). Most field data were collected along two composite stratigraphic sections (the Kallipetra and Sfikia sections, Fig. 4).
The base of the basin is exposed close to the contact with the Pelagonian
basement. Locally, the latter consists of a thick package of white, foliated
cataclasite containing both serpentinite and gneiss fragments (Fig. 3).
Directly overlying the cataclasite is a very dark massive fractured to
brecciated serpentinite, followed by pebbly sandstones and well-bedded dark
gray limestones. Elsewhere, the base of the basin is characterized by a
thick package of serpentinite-rich conglomerates, breccias, and minor
amounts of dark gray limestone (Fig. 4a). The conglomerate is
clast-supported and poorly sorted, with a dominance of sub-rounded to
rounded clasts greater than 15 cm. The conglomerate is composed of dark
green to black serpentinite (
Calc arenites are observed throughout the basin, typically at intermediate stratigraphic levels (Fig. 4b). The arenites range from fine- to coarse-grained, are medium to thickly bedded, and often display slumping folds. Locally, these folds and the synsedimentary gravity faults show a top-to-the-NE vergence. The quartz content varies with location, with the highest proportion of quartz being in the northwest region of the study area. Locally, the calc arenites consist of medium- to coarse-grained poorly sorted pebbly sandstones with 1–6 cm sized clasts of red arenite and red–pink carbonate. Very distinctive thinly bedded and laminated red and green marly limestones occur at intermediate-to-high stratigraphic levels (Fig. 4b). The red layers range from 2–5 cm thick, and green layers typically range from 0.5–2 cm thick. These deposits represent the deepest pelagic facies of the basin.
Towards the top of the basin, massively bedded conglomerates and breccias are often interbedded with the calc arenites and pebbly sandstones and consist of limestone, bioclastic limestone, arenite, marl, serpentinite, mudstone, and calcareous mudstone as rounded to sub-rounded clasts in a calcareous matrix.
Lateral variations in facies occur frequently, the most evident being the changes observed from the northwestern to the central and southeastern portions of the mapping area. In the northwest (Fig. 4b), the stratigraphy is dominated by coarse to pebbly sandstones, breccias, and conglomerates, whereas shaley limestones, marls, and mudstones prevail in the southeast (Fig. 4a). In the northwest, lithic fragments of quartz, gneiss, and marble are major components of the coarse sediments, with quartz content ranging from 45 % at the base to 90 % up-section (310 m, Fig. 4b), where serpentinite forms a minor component. In addition, olistoliths and evidence of slumping are frequent at high stratigraphic levels in the northwestern sector (Fig. 4b). This differs greatly from the southeastern sector (Fig. 4a), whereby slumped calc arenites with olistoliths appear only at the base of the section and the average quartz content is lower.
The top of the Kallipetra Basin is marked by the occurrence of rudist mounds, five of which, some tens of meters thick, were identified in the study area. The mounds produce prominent cliffs and dome-like structures in the topography. Each mound can be separated into four different facies associations (Fig. 5): (i) the serpentinite and Kallipetra carbonate breccia (SKB); (ii) the mound core; (iii) mound flank; and (iv) the mound top.
Stratigraphic sections of the Vardar ophiolitic complex, mound top, and mound flank.
The SKB is a sub-angular, moderately sorted, clast-supported breccia that is poorly bedded and massive. Clasts comprise serpentinite, dark gray limestone, rudist-rich microsparite, pink micrite, and minor lithic fragments like quartz, feldspar, and some dark pyroxenes. The rudist microsparite and pink micrite clasts are identical to the mound core. The matrix is composed of a fine- to medium-grained calcareous arenite. Orbitolinids were discovered in a clast by Schenker (2013). The SKB is usually found on the southern side or stratigraphically below the mound.
The mound core is characterized by light gray to pink, massively bedded
micrite and microsparite, in which float abundant whole rudists.
The mound flank is a heterogeneous lithology that varies with distance from
the mound core and location within the basin. In general, a moderately
sorted, clast-supported breccia containing large, angular clasts of rudists,
red limestone, greenish marls, micrite, and minor serpentinite clasts occurs
closest to and on the northern side of the mound core. The number of clasts
decreases into a matrix-supported breccia with a marly, green-colored
matrix. The serpentinite content gradually increases up-section, and
gravel sandstones contain
Stacking of serpentinite-rich breccias always occurs on the southern slope of the rudist mounds. Flank deposits, either marls or a succession of sandstones and breccias, dip away from the mound core always on the northern mound side.
The mound top, where observed fully, is approximately 6 m thick and stratigraphically overlies the mound core (Fig. 5). It generally consists of several meters of very poorly sorted, angular to sub-angular gravel of serpentinite and quartz within a white calcareous matrix. A thin layer of rudist-rich, elongated carbonate clasts overlies the gravel. There is a gradual transition into a clast-supported conglomerate with a reddish calcareous matrix, plus arenite and minor serpentinite clasts. The clasts of this conglomerate are very deformed, where the VOC tectonically overlies them. The full stratigraphy of the mound top was only observed at one, the southernmost, rudist mound (Fig. 5).
Although significant amounts of sample were collected for biostratigraphic analysis, nearly all of them were barren, or included dissolved, silicified, or recrystallized nannoplankton and foraminifera, making most species indistinguishable. Table 1 summarizes the recognizable planktonic foraminifera that were only found near the northernmost mound (Asomata Quarry).
Table of observed foraminifera.
The orbitolinid found in sample M2-TS3,
Species abundance and totals of calcareous nannofossil were
semi-quantitatively evaluated as F (frequent) and R (rare). In the
studied section (M2), the major calcareous nannofossil events in
stratigraphic order are as follows: the first occurrence of
In the marls and marly limestones of the Kallipetra Basin, the foliation is
mostly parallel to the bedding and defined by flat and elongated quartz
clasts and clay minerals. Bedding and foliation dip at a very low angle
either to the NW or to the SE due to bending around an axis plunging
shallowly to the NE (Fig. 6a, b; Table S3 in the Supplement). Stretching lineations are
observed mostly on lamination surfaces in fine-grained marls, limestones, and
mudstones, and they are formed by the alignment of elongated clay minerals.
Mineral lineations occur mainly in the Pelagonian basement where the long
axis of elongated quartz and feldspar crystals are aligned. In all the
lithologies the lineations strike NE–SW at low dip angles (
Lower hemisphere stereoplots of:
The top of the Kallipetra Basin is tectonically covered by the VOC. In the
profiles A-A
Geologic cross sections of the mapped area, colors corresponding to those on the geological map (Fig. 3).
Late normal faults trending W–E to NW–SE and two transtensional to
strike-slip faults crosscut the Pelagonian basement, the Kallipetra Basin, and the VOC (Fig. 3). Most steep normal faults plunge to the NE (Fig. 8).
Low-angle fault zones are observed within the VOC, dipping approximately
35
Thirty-seven samples for illite crystallinity analysis were taken up-section in the
north part of the Kallipetra Basin (Kallipetra section). Four samples were
unsuitable for illite crystallinity analysis as the
Illite crystallinity data.
Stratigraphically higher samples have a KI ranging from 0.09 to 0.25, and stratigraphically lower samples have a KI of 0.39. The KI appears to increase down-section for samples containing only non-detrital illite. The sample with the lowest KI of 0.091 is characterized by an XRD pattern that reveals the presence of chlorite. The sample with a KI of 0.14 contains paragonite which indicates epizone conditions.
The samples containing detrital illite are limited to stratigraphic heights
between 300 and 350 ms and show a large range of KI between 0.14 to 0.383.
Non-detrital illite, on the contrary, is mostly confined to stratigraphic
heights above 400 m. Using the anchizone limits as calibrated for our lab
(see Sect. 3.2; Jaboyedoff et al.,
2000) and given the fact that the effects of detrital micas disappear in
the anchizone (
We collected our samples along a down-section direction within the
Kallipetra Basin: the only ones that produced enough countable zircons are
from close to the contact with the VOC. Results are reported in Table 3. The
two successful samples are from the same location but from two different
layers: a sandstone and a conglomerate. Both rocks are sheared and contain
newly formed chlorite (Fig. 9). In sample V1504, 61 grains could be counted
on the 10.5 and 17.5 h etch. In sample V1503, 79 grains could be counted
on four mounts with the three different etch times (10.5, 14, and 17.5 h).
Both samples consist of multiple age populations as attested by the
Zircon fission track ages of samples taken in the mapped area.
Zircon fission-track data. Variable amounts of zircons were
analyzed on multiple mounts for each sample that were etched for different
times. As a fluence monitor, a glass standard CN1 with a
The orthogneiss- and serpentinite-rich composition of the previously
described basal cataclasite suggests that it was formed prior to or at the
same time as the formation of the Kallipetra Basin and mainly at the
expense of the Pelagonian basement and of the VOC. These normal faults
crosscut duplicates of Pelagonian mylonitic marble and must be younger than
ca. 120 Ma (Schenker et al., 2014). Normal faulting during or
following the exhumation and doming of the Pelagonian zone from the late
Early Cretaceous (Schenker et al., 2014;
Schenker et al., 2015) probably contributed to subsiding the deformed wedge
below sea level to create the basin. As discussed in Sect. 2.1, the onset
of the Cretaceous basins along the Pelagonian zone is diachronous and
remains enigmatic. However, some evidence may suggest that the Albian–Aptian
topographic response to the Early Cretaceous continental accretion was
uneven along strike of the Pelagonian zone, likely due to northward-decreasing shortening to a wider Rhodope to the south, thinning out to the
north (Fig. 1). Indeed, to the south, non-metamorphosed Pelagonian sediments
showing a sedimentary Aptian–Albian hiatus (
Serpentinite-rich conglomerates represent the first sediments deposited within the Kallipetra Basin through subaerial erosion of the VOC, which created an uneven topography. North of the study area, conglomerates containing serpentinite pebbles and Pelagonian marbles of Albian age are observed on the eastern Pelagonian zone border, demonstrating prior deep erosion of the obducted ophiolitic sheet likely of the Pelagonian Zone (Mercier and Vergely, 2002). The occurrence of conglomerates followed by a succession of calc arenites at the base of the basin indicates shallow marine depths. Marble olistoliths and slumping at the base of the Sfikia section (Fig. 4a) indicate instability during the first phases of basin formation and the presence of a proximal steep slope in which gravitational instability drove slumping. In the northwestern part of our study area, the presence of orthogneiss blocks, the dominance of quartz, feldspar, gneiss, and marble lithics in the sediments, and the lack of such components in the southeast suggest an intra-basinal high, emergent land or continent existed northwest of the Kallipetra Basin, where the Pelagonian basement was exposed. In the southeastern part of the study area, the dominance of silty limestones, marls, and lime mudstones and the rare presence of olistoliths indicate that there was a deepening of the basin towards the southeast. The mid part of the Kallipetra Formation is devoid of serpentinite coarse detritus, suggesting the initial fault scarps were smoothed by sediments. This expansion of the basin toward the southern slope is documented by the transgression of Kallipetra deposits onto the VOC, recorded in the study zone. It is worth noting that this basin widening corresponds to the deepest facies in the basin, likely correlating the Cenomanian–Turonian eustatic sea level high (e.g., Haq, 2014). Mercier (1968) and Sharp and Robertson (2006) also record marine transgression and eastward deepening of mixed carbonate–clastic successions of the combined Pelagonian and western Almopias zones.
Rudists constituted more than 60 % of reef frames during the Aptian and
Albian and became the most dominant frame-building organism in the Late
Cretaceous (Scott, 1988; Voigt et al., 1999).
Widespread tectonic extension combined with eustatic continental flooding
occurring around the Cretaceous Tethyan Ocean allowed the growth of broad
carbonate platform complexes, on slopes from a few up to 40
The mound flanks consist of a succession of sandstones, breccias, and occasionally marls. Our observation of polymictic breccias on the south–southwestern mound slopes bears an important paleogeographic meaning. In this case, the presence of serpentinite clasts indicates the breccias were not formed solely from the erosion and collapse of mound flanks, but rather they were derived from an ophiolitic source upslope from the mounds, possibly associated with new fault scarps in the south–southwestern slope of the Kallipetra Basin. On the northern flanks of the mounds, the sediments display a shallower dip and are interfingered with the mound talus breccias. The youngest and northernmost mound at Asomata Quarry displays pelagic marls and limestones at the northwestern flank, suggesting deeper bottom conditions to the N and NE. On the other hand, the absence of serpentinite detritus in the mound flanks other than the southern ones documents a shadow effect of the mounds with respect to the south–southwestern provenance of the serpentinite clasts. This evidence corroborates the presence of a slope dipping to the north–northeast. North of the study area, Cenomanian-aged hippuritid-bearing rudist mounds have been observed, where Sharp and Robertson (2006) suggest they also developed on an east-facing ramp. These authors also observe younger Campanian–Maastrichtian rudist biostromes that developed on an east-facing ramp–shelf margin in the Pelagonian and western Almopias zones. However, as discussed in the following sections, these must have formed subsequent to Kallipetra Basin closure, likely in basinal areas not involved in Turonian compression. The increasing serpentinite content in the sandstones and breccias up-section suggests that the ophiolitic source was moving closer to the mound structure and gradually providing material to the slope. The positioning of the flank deposits and the northeastward directed stacking pattern of the two or three youngest mounds, with the highest – and therefore youngest – mound being at the Asomata Quarry in the northeast of the study area, suggest that in the upper part of the Kallipetra stratigraphy, there was a movement of the ophiolite (VOC) from SW to NE providing at first the slope for the growth of the mounds and then the burial for them (Figs. 12 and 13).
The Kallipetra Formation was deposited on top of the eroded Pelagonian
continent and obducted ophiolite following the collision-related burial and
cooling or exhumation of the Pelagonian zone at
The red and green limestones can be loosely correlated across the Kallipetra
Basin and they first occur at approximately
Rudist mounds and breccias are lacking in the Kallipetra sediments found on top of the VOC (Figs. 3 and 8), whereas fine sediments dominate. The fine material suggests that the tip of the VOC was under sea level, with a transgressive trend and widening of the basin, allowing the onlap of fine material over the VOC slopes at the same time as red and green marl deposition. During this time, the source area for sediments is moving away. This Kallipetra material overlying the VOC is somewhat separated from the main Kallipetra Basin sediments, possibly through a structural high or as perched basins (Fig. 13). Alternatively, these deposits represent one flank of the basin that was subsequently tectonically emplaced over the basin, suggesting this movement was just a few kilometers and a local event.
The stratigraphic thickness between the red and green marls, and the marls
adjacent to the mound core containing the
Our new ZFT samples come from the top of the Kallipetra Basin where the
depositional age should not be older than 92 Ma and therefore should be
younger than the ZFT central ages of our samples, which range between 156
and 177 Ma (Fig. 10). However, both ZFT samples contain a few young grains
with ages overlapping with the depositional age of the Kallipetra Formation
(Fig. 10). They are located a few hundred meters to the south of a previous
sample, 10-029, that is adjacent to the contact with the VOC (Schenker et
al., 2015; Fig. 10). The age range of this sample is from 52 to 340 Ma, and
it consists of 16 grains that define only one age population centered at
Two more samples were previously dated along the Kallipetra section where we
collected our new illite crystallinity data (Fig. 10; Schenker et al.,
2015). There, 20 grains from the sample at the top (10-128) of the section
define an age range between 39 and 102 Ma and a central age of
The different ZFT age ranges and central ages hint at highly variable degrees of annealing. Our petrographic and illite crystallinity data constrain a strong, inverse, vertical (up-section) thermal gradient, but they cannot discriminate between possible lateral gradients across the basin. However, they indicate that the Kallipetra Basin has been subject to temperatures that locally could have totally or partially annealed our samples. Whether these gradients are reflected by the ZFT central ages or grain-age distributions must be carefully weighed against other factors that could also affect our results. In fact, we processed the new and the old samples purposefully in different ways because, while processing the previous set of samples, we realized that the low number of available zircons limited the applicable etch procedure, which was not optimal to reveal the full age spectra of our samples. However, even though at the time we opted for an etch procedure aimed at maximizing the young grain ages, our results indicated that the annealing degree of our samples might have been incomplete. With the new samples, we aimed at verifying the degree of annealing by maximizing the zircon yield that allowed applying a multiple etch procedure. This in turn revealed that in fact there are wide age distributions in the new samples, which include non-reset ages, and this confirmed our previous observations on a partial degree of annealing. Unfortunately, our new data do not answer all the questions concerning the ZFT ages in the study area but highlight a complex thermal and annealing record.
The KI data constrain an inverse geothermal gradient at the top of the
Kallipetra Basin from
Schematic diagram showing the inverse geothermal gradient at the contact between the VOC and Kallipetra Basin.
The illitization reaction (i.e., the conversion of smectite-rich I-S into
illite-rich I-S) is also dependent on the availability of K
Overall, our data document an inverse thermal gradient of the Kallipetra Basin, pointing to a syntectonic heating event that produced a transient, inverse, non-linear, and disturbed geotherm (Fig. 11). The sedimentary history suggests that the closure of the Kallipetra Basin by the VOC occurred just after the deposition of the ophiolitic debris that buried the rudist mounds, when the sediments were porous, permeable, and saturated. Although the ultimate sources of this heat have not been established, the non-reset to partially reset FT ages testify that this syntectonic heating event formed in the Late Cretaceous, during the closure of the basin in the Turonian. Cooling slightly postdates the deformation as the youngest ZFT population is older than the Turonian closure.
The stacking pattern of the rudist mounds documents closure of the Kallipetra Basin through a NE-facing slope and serpentinite detritus supplied from the SW. This pattern cannot be explained by the activity of a normal fault (Fig. 12-2a), since, in this case, the stacking of the mounds should have been to the SW, following the widening of the basin. Therefore, only a NE-verging thrust of the VOC found at the southwestern margin of the basin can explain the observed rudist mound stacking pattern, as well as the serpentinite breccias and the northeast mound shadow (Figs. 12-2b and 13). Kinematic indicators, such as shear bands, stepover structures, and sigma clasts along the upper tectonic contact of the Kallipetra Basin, indicate a top-to-the-NE tectonic movement. If the NE-dipping contact of the VOC over the Kallipetra Basin is regarded as the pristine fault orientation, the sealing may be interpreted as a Turonian or younger normal fault (Fig. 12-2a.i). However, this interpretation fails to explain the observed cutoff angles, since the tectonic contact should have cut the basinal deposits down-section, and this is not the case (Figs. 8 and 12-2a.ii). The cutoff of the Kallipetra deposits is compatible only with a NE-directed thrust (Fig. 12-2b.i). Post-Turonian tectonics are considered responsible for the northeastward block rotation and the normal faulting (Fig. 12-2b.ii).
A series of sketches to demonstrate the opening of the Kallipetra Basin in the Cenomanian and the closure of the Kallipetra Basin under normal faulting or thrust faulting conditions. Only a NE-verging thrust of the VOC found at the southwestern margin of the basin can explain the observed rudist mound stacking pattern, as well as the serpentinite breccias and the northeast mound shadow. Colors and patterns correspond to those on the geological map (Fig. 3).
Schematic diagram showing the sedimentary and tectonic environment of the Kallipetra Basin during the Turonian and the overriding of rudist mounds by the resumed thrusting of the VOC. Colors correspond to those on the geological map (Fig. 3).
The partially to fully annealed ZFT ages combined with illite crystallinity and crystallization of syntectonic chlorite, indicating high temperatures at the tectonic contact, but only deep diagenetic conditions below, suggest that the tectonic movement occurred in the Turonian when sediments were not fully compacted and still permeable. Therefore, the Kallipetra Basin was sealed in the Turonian by SW to NE tectonic transport of the VOC (Figs. 12 and 13). Northeast-directed thrusting is documented some 50 km north of our study area in the Almopias zone by Vergely and Mercier (2000), although it is considered Tertiary in age by the quoted authors. However, our results differ from previous studies that have documented progressive deepening from the Aptian up to flysch-like Maastrichtian to Paleogene sediments (Mercier and Vergely, 2002) or the development of thick Santonian–Campanian carbonates followed by a foredeep succession indicating dramatic subsidence in the Late Maastrichtian (Sharp and Robertson, 2006). The observed stratigraphy from this study, such as olistoliths, breccias, and slump deposits, suggests deposition very close to steep basin margins. Therefore, it is plausible that the closure of the Kallipetra Basin observed here only affected the margins of a larger-scale basin, which experienced continuous deepening and transgression as observed to the north of the study area by Mercier and Vergely (2002). Alternatively, the Kallipetra Basin may not be laterally continuous, and its birth and closure would have only had local significance.
During the Late Cretaceous to Eocene, along the eastern Pelagonian margin, the dominant deformation at the regional scale is SW-verging thrusting (e.g., Schenker et al., 2015, and references therein). Bivergent thrusting occurred locally but later in time during the late Late Cretaceous or Tertiary (Vergely and Mercier, 2000; Brown and Robertson, 2003; Katrivanos et al., 2013). Thus, the sealing of the Kallipetra Basin occurred earlier than or in the very early phase of this regional deformation event, although the direction of tectonic transport of the VOC above this basin is opposite to the common SW vergence of thrusting. This apparent difference may be explained by a localized basin inversion rather than a regional tectonic event that predated the start of the regional convergence in the Late Cretaceous, i.e., the Campanian at the earliest (Aubouin, 1973; Baumgartner, 1985; Godfriaux and Ricou, 1991; Bonneau et al., 1994; Papanikolaou, 1997; Brown and Robertson, 2003; Grubić et al., 2009; Ustaszewski et al., 2009; Kilias et al., 2010; Katrivanos et al., 2013; Schmid et al., 2020).
Overall, the thermo-tectono-sedimentary history documents a basin that likely formed in the early Cenomanian over a suture with accreted Jurassic ophiolites in the hanging-wall and an Aptian metamorphic basement in the footwall. This attests its intracontinental position within an orogenic wedge in a tectonic scenario similar to the model proposed by Schenker et al. (2014) (Fig. 2). The Kallipetra Basin cuts and covers the Pelagonian basement with an Aptian regional metamorphism that cannot be associated with the metamorphic sole beneath the Jurassic–Early Cretaceous obduction but rather with a Barrovian metamorphism formed during the collision of the Pelagonian zone with the Rhodope (Schenker et al., 2014). Accordingly, geodynamic models (cf. Fig. 2) that invoke a Jurassic–Early Cretaceous obduction without an Aptian crustal-scale accretion (e.g., Sharp and Roberson, 2006; Froitzheim et al., 2014) or the distal position from the Pelagonian zone of the Rhodope during the Early Cretaceous (e.g., Ricou et al., 1998; Froitzheim et al., 2014) do not agree with the thermo-tectonic history of the Pelagonian zone prior to the formation of the Kallipetra Basin. The extensional phase that opened the Kallipetra Basin remains enigmatic and may be associated with an isostatic re-equilibration of the orogenic wedge or with a far-field plate tectonic reorganization (e.g., Matthews et al., 2012). The closure of the basin anticipates the beginning of resumed ophiolitic imbrication in this sector of the Internal Hellenides in the Turonian and is considered a local basin inversion. If the actual Hellenic subduction is considered active since at least the Early Cretaceous (van Hinsbergen et al., 2005), the closure of the Kallipetra Basin could be seen as early evidence in the upper crust of the initiation of the Hellenic slab.
The evolution of the Kallipetra Basin documents the transition from
extension to compression during the early Late Cretaceous along the eastern
margin of the Pelagonian zone in northern Greece. The history of the
Kallipetra Basin can be summarized as follows:
The sediments of the Kallipetra Basin were deposited between the early
Cenomanian ( As the basin widened, a topographic high located to the NW and exposing
Pelagonian basement rocks became the main source of siliciclastic detritus
to the basin. The basin widened and deepened to the point when no clastic
input reached it. This time might correlate with the global
Cenomanian–Turonian sea level transgression. Carbonate sediments were
produced by pelagic organisms and by rudist-rich microbial mounds growing on
the southwestern slopes of the basin (Fig. 13). The terrigenous input was later renewed, and the main source were ophiolitic
rocks to the south or southwest, which provided breccias stacking up
against the southern flanks of the rudist mounds. The progressive increase in detrital input restricted the environments of the rudist mounds.
The ophiolitic rocks overrode the Kallipetra Basin from the SW, causing
uneven deformation of its sediments (Fig. 12). Thrusting was associated with
a high inverted geothermal gradient that caused illitization,
crystallization of chlorite, and partial-to-total annealing of the fission
tracks in detrital zircons close to and increasing towards the top of the
basin. Deformation, illitization, and zircon fission track annealing occurred
during the Turonian and were followed by cooling in the late Late
Cretaceous, anticipating the beginning of the resumed tectonics in this
sector of the Internal Hellenides by about 10 Ma.
Additional data are available from the corresponding author upon request.
The supplement related to this article is available online at:
LRB is the primary author, conducted the study for her MSc thesis at ETH Zurich, and wrote the paper. FLS conceptualized the original research goals and aims of this study, and contributed to the interpretations, introduction, and background sections of this paper. MGF conducted the fission track analysis, provided contributions to Sects. 1, 3.3, and 4.5 and to interpretations of data and also provided mentorship. MC prepared samples for and conducted all planktonic foraminifera analysis for this study. TA prepared samples for illite crystallinity analysis and conducted the latter as well as giving constructive suggestions on the final paper draft. VP contributed to the interpretation and provided supervision and mentorship in all aspects of this study.
The authors declare that they have no conflict of interest.
This article is part of the special issue “Inversion tectonics – 30 years later”. It is not associated with a conference.
We thank the two anonymous referees for constructive reviews that significantly improved the paper. A special thanks goes to Tiemen Gordijn for field assistance in the Kallipetra Basin and to the welcoming community of Sfikia for hosting us. We thank Annie Arnaud-Vanneau and Daniel Bernoulli for help with fossil identification.
This paper was edited by Jonas Kley and reviewed by two anonymous referees.