SUPPLEMENTARY SECTION 12,800 Years Ago, Hellas and the World on Fire and Flood Volker Joerg Dietrich, Evangelos Lagios and Gregor Zographos

SUPPLEMENTARY SECTION 12,800 years ago, Hellas and the World on Fire and Flood Volker Joerg Dietrich, Evangelos Lagios and Gregor Zographos

Supplements

1 The Geotectonic Framework of the Pagasitic Gulf 1.1 Alpine Tectonic Structures 2 Surficial Cataclastic and Brittle Deformation 2.1 Macroscopic Scale (Breccia Outcrops) 2.1.1 Striation and Shatter Cones 2.2 Microscopic Scale 2.2.1 Planer Deformation in Quartz 2.2.2 Planer Deformation in Calcite 2.3 Metamorphic and Post-Alpine Hydrothermal Activity (Veining) 3 Geophysical Investigations of Pagasitic Gulf and Surrounding Areas Gravity Measurements and Modelling

1 The Geotectonic Framework of the Pagasitic Gulf

The Geotectonic frame of the Pagasitic Gulf is best exposed in the sickle shaped Pelion Peninsula (Figs. 1&2) and applies to all mountain ranges and coastal areas around the gulf, which are part of the “Internal Alpine-Dinaride-Hellenide Orogen”.

Fig. 1 Google Earth image of the Pagasitic Gulf – Mt. Pelio area; bathymetry according to Perissoratis et al. 1991; Korres et al. 2011; Petihakis et al. 2012. White Circle on the western side of the image: The Zerelia Twin-Lakes: Two Possible Meteorite Craters (Dietrich et al. 2017). 0 1.1 Alpine Tectonic Structures

The internal structure of Pindos and Pelagonian thrust sheet units is extremely complex and has not yet been worked out in detail. In addition, towards north overthrust units of the Axios-Vardar realm cover the Pelagonian thrust sheets (Fig. 2).

Fig. 2 Synthetic cross section through the Olympos region between the “External Hellenides” and the “Axios/Vardar tectonic nappe system” after Schenker et al. (2014). Moho below the Hellenides from Sachpazi et al. (2007), below the Vardar/Axios zone Papazachos (1998).

Besides mappable overthrust boundaries, intense internal folding has only been observed. Both units underwent a post-Cretaceous and pre-Eocene high-pressure/low temperature deformation and recrystallization age (Ferrière 1982; Katsikatsos et al. 1986; Schermer 1990), followed by retrograde greenschist metamorphism and ductile deformation between Eocene and Miocene time (Schermer et al. 1990; Schermer 1993; Lips et al. 1999; Nance 2010; Schenker et al. 2015), which resulted in a large scale anticlinal structure with a general NW-SE axis, doming of the Pelio mountain and dip-slip movement toward the present-day east. Ductile deformation may have terminated at Late Miocene. During Plio-Pleistocene uplift and brittle normal faulting continued and extended through Pleistocene to recent being responsible mainly for the NE-SW faults (Fig. 3). All pre-Neogene rocks are metamorphosed and have been grouped into two major tectonic units, a lower part belonging to the Pelagonian nappes and a higher part consisting of dismembered Maliac - Eohellenic thrust sheets (Wallbrecher 1976; 1977, 1983; Ferrière 1976, 1977, 1982; Jacobshagen 1986; Jacobshagen et al. 1979; Robertson 2002; Gartzos et al. 2008; Schenker et al. 2015; Ferrière et al. 2017; Nirta et al. 2018). Both units are overthrusted by Upper Cretaceous – Palaeocene breccias, limestones, flysch (Plesidi, Lechonia), conglomerates and ophiolitic remnants as

1 olistolitic bodies (Fig. 3). Minor outcrops of Neogene rocks occur only along the eastern coast east of Liri and south of Paltsi.

Fig. 3 Simplified tectonic map of the Pagasitic Gulf – Mt. Pelio area, showing the spatial distribution of the Pelagonian thrust sheet units; compiled after the Geological Map of Greece 1:50000, Sheets: Almyros (Marinos et al. 1962); Volos (Katsikatsos et al. 1986); Zagora-Siky (Katsikatsos et al. 1987); Argalasti (Katsikatsos et al. 1989) and special publications (Wallbrecher 1976,1977,1983), Ferrière 1976,1977,1982; Jacobshagen 1986; Jacobshagen et al. 1979. Basis: Google Earth.

Post-orogenic extension followed the Alpine collision in the Aegean microplate starting during Miocene and ongoing today. Consequently uplift, exhumation and erosion took place, e.g. in excess of 10km of exhumation for the Olympus, Ossa, Pelio mountain range (Kilias et al. 2002; Nance 2010). An early ductile deformation phase accompanied with mylonitization was postulated in the Olympus realm by Schermer et. al. (1990) and Nance (2010), which started at Early Miocene ca. 23-16 Ma. (U-Th)/He dating combined with apatite fission-track from Pelagonian unit rocks in the region around Mount Olympus revealed a rapid cooling event at Middle Miocene (14-10 Ma; Walsh 2013; Schenker et al. 2015), which proofs rapid uplift processes.

Fig. 4 Agria Bay and town of Agria (center) with Mt. Pelio 1550 m; the Pelagonian nappe pile: in the foreground the Triassic Jurassic marble formations on top of the Upper Paleozoic to Triassic schists (background); from left to right: Kastri 429m, Kteni, 524m, Kardaras 296m and Efzonos 315m; in the background the Upper Paleozoic to Lower Jurassic Pelagonian tectonic unit. Photo V. Dietrich.

Fig. 5 Overview southern part of Pagasitic Gulf and Trikeri street: view from Miriovriti outlook towards southwest; in the background (right) the Orthris Mountain range. In the foreground smooth surface of the Eohellenic thrust sheets of uncertain Triassic to Jurassic age. Photo V. Dietrich.

Towards south, the northern Upper Paleozoic to Lower Jurassic Pelagonian tectonic unit is overthrusted by an approximately 700m thick unit of Eohellenic thrust sheets of uncertain Triassic to Jurassic age, which are composed of epidote-actinolite- chlorite schists, mica schists, phyllites, crystalline limestones and platy marbles, as well as major metamorphosed mafic bodies and small lenses of highly tetanized serpentinites (Figs. 3,4). In hills along the east coast of the Pelio Peninsula between Lambinou and Aghios Dimitrios light grey to greenish schistose gneisses and platy gneisses with a thickness up to 350 m, in places with intercalations and lenses of grey crystalline limestones and meta-basalts overlay the Triassic-Jurassic marbles. South of Kalamaki, this formation is quarried as the famous Pelio stone (Figs. 5-7). Metamorphic conditions within this unit vary from greenschist to glaucophane schist facies. The lower tectonic thrust sheet has been designated as Trikeri unit and the overthrusted higher thrust sheet as Argalasti unit in the Geological Map of Greece. 3

Fig. 6 Pelio-Stone quarries in the walls of a NE-SW valley north of Neochori within the up to 700m thick unit of Eohellenic thrust sheets of uncertain Triassic to Jurassic age. White Triassic to Lower Jurassic marbles (left side), overlain by greenish gneisses with marble intercalations, possibly Jurassic in age, are overthrusted by brownish mica schists and phyllites. Photo V. Dietrich.

Fig. 7 Close up of the upper southwestern part of NE-SW valley north of Neochori (Fig. 6). Clearly visible is the slightly folded but sharp overthrust of the upper thrust sheets, made up of mica schists, greenschists and phyllites with intercalations of crystalline limestones. Photo V. Dietrich.

Fig. 8 View from road at Metochi village towards southwest, Dhiasela, Trikeri Island, and Kokkinovrachos Mt. (background, center); to the left the villages of Horto and Milina; note that all slopes of Mt. Tiseo dipping toward the Pagasitic Gulf. Photo V. Dietrich.

In addition, highly tetanized ophiolitic remnants (serpentinites, meta-gabbros, diabase, meta-basalts) occur as “klippen” in the areas of Argalasti, St. George and on old trickery island (Fig. 8), which are interpreted as remnants of oceanic crust of a former Maliac ocean (Ferrière et al. 2017). All thrust sheets are overlain with tectonic contacts less metamorphosed Lower Cretaceous schists, phyllites, metasandstones and small ophiolitic “olistolites”, which turn into Upper Cretaceous crystalline limestones with brecciated conglomerates at their base.

1.2 Neogene to Quaternary Faults and Brittle Tectonics

Strong and rapid uplift since Upper Tertiary in the Olympus, Ossa, Pelio mountain range, accompanied with tensional tectonics throughout the Internal Hellenides, may have caused deformation at low temperatures and at all scales: microscopic foliation and lineation, cataclastic and brittle deformation, mesoscopic fractures and macroscopic faults. These features have been recognized by many geologists mapping the region as unique phenomena (Wallbrecher 1976; Ferrière 1982; Katsikatos et al. 1986, 1989; Galanakis et al. 1998; Papanikolaou and Migiros 2008). During Plio-Pleistocene normal faulting generated the NNW-trending fault systems west and east of the Olympus, Ossa, Pelio mountain range, which were designated by Galanakis et al. (1998) in the southern part as Pagasitic and Pelio zones. Strong Quaternary uplift has been recognized along the east coast of the Pelio peninsula (Stiros et al. 1994) based on notches, relict beach terraces with clastic sediments and AMS radiocarbon dating of fossils in several localities, e.g. sites of Damouchari, Mylopotamos, Labinou, Paltzi etc. The formation of NE-SW and ENE- WSW fault are regarded as result of extensional crustal movements during Pleistocene and Holocene (Fig. 8).

Fig. 8 Neogene and Quaternary fault pattern combined with Plio-Pleistocene and Holocene lacustrine and fluviatile deposits in the Almiros plain and Pagasitic Gulf area derived from aerial photographs and mapping (Caputo and Pavlides 1993; Galanakis et al. 1998), redrawn on the Google Earth shaded 3D image.

Three sets of faults have been recognized in the geotectonic frame of the Pagasitic Gulf (Fig. 8; after Ferrière 1982; Galanakis et al. 1998; Kreemer et al. 2004; Papanikolaou and Migiros 2008): (a) The approximately E-W trending (mostly normal) faults, represented in the northern part of the Almyros plain and N. Anchialos fault zone, as well as in the central and southern part of the Almyros plain (Almyros zone, Galanakis et al. 1998) may have been active since Early Pleistocene, which generated a shallow basin structure. (b) The common NE-SW trending faults (N 030° ± 10° and N 050° ± 10°; e.g. Trikeri zone), which were interpreted as strike slip faults (Ferrière 1982) parallel to the south marginal fault of the North Aegean Basin, and thus an extension of the North Anatolian Fault. (c) A common third set of common faults, NW-SE to NNW-SSE (N 320° ± 5°; e.g. Pagasitic zone) accompanied with down-faulting movements occur towards the lowland of Volos, shoreline and offshore areas of Pagasitic Gulf, as well as towards their intersection points with the transverse E-W trending oblique normal faults, and seemed to have formed under extensional movements, since they are associated with many springs. This structural configuration Plio-Pleistocene faulting could have led to a basin structure a prerequisite of large shallow lake (Galanakis et al. 1998). A large about 10 km offshore NW-SE trending normal fault runs parallel to the shoreline and forms the western boundary of the North Aegean Basin (Papanikolaou et al. 2006).

2 Surficial Cataclastic and Brittle Deformation

Evidence of a meteoric impact is normally given by pressure and temperature effects on the underground of the impact site, e.g. on rocks and soil inside the impact crater, the distal environment and the occurrence shock-infected or molten ejecta. In the case of a possible impact forming the Pagasitic Gulf, no material from sea floor is available. Besides its shallow bowl shape and a gravity anomaly, only the distal environment has remained for investigation. In fact, along the embankments and under the young Quaternary surface fracturing, brecciation, cataclastic and brittle deformation is present in different lithologies along the shores (localities in Fig. 9). A large variety of breccia occur, in most cases monomict breccia, clast and matrix supported with clasts from millimeter to decimeter size. However, typical impact melt breccia (matrix-melt breccia) and suevite (breccia with glass, crystal and lithic fragments) as well as distal ejecta layers containing spherules have not been recognized. The different types of breccia are shown in the following figures 10 to 27.

2.1 Macroscopic Scale (Breccia Outcrops)

Fig. 9 Morphotectonic map with localities of major characteristic breccia (red triangles) and cataclastic deformation around Pagasitic Gulf; Numbers refer to figures in the text; a compilation of rotational landslides (surface slumping and listric faults in yellow) with radial normal and strike/slip faults in red), all dipping towards the gulf; strike and dip signs were taken from map sheets (1: 50

7 000 Geological Map of Greece, Katsikatsos et al. 1978, 1987, 1989; Marinos et al. 1957, 1962); red triangles are locations of major breccia occurrence; drawn on Google Earth image.

The outcrops along the northwestern coast between Nea Anchialos, Cape Agistri and the hill of Amphanae (Gulf of Volos) and in the main road cuts expose schists, gneisses and crystalline limestones, which are cut by numerous faults (Figs. 10a-e). In general, all lithologies dip towards the gulf. Slumping appears using the 3D observation in the Google Earth maps. Most of the deformation, gentle folding and schistosity can be assigned to pre-Quaternary orogenic processes. Post-deformational features are steep normal down faulting parallel to the coastline (Anchialos fault zone) as well as minor perpendicular fault systems. The gneisses are deeply fractured in several areas (e.g. in the quarry below the hill of Amphanae, Figs. 10e&f), the surface in addition chattered. Between Kritharia and Stavros crystalline limestones and schists are distorted and brecciated, partly filled with reddish Pleistocene red beds and hydrothermal deposits, dipping towards sea. A similar situation of intense fracturing is evident along the shore of Cape Ghoritsa for almost two kilometers between Volos and the cement factory of Asteria/Agria (Figs. 11,12). Outcrops are best exposed in the cliffs and in the main road cut. The coastal cliffs demonstrate that fracturing and brecciation in the middle Triassic to Jurassic marbles is a natural phenomenon and not the result of blasting. However, the latter effect of brecciation due to strong blasting in some cases along the road cut cannot be excluded. The construction of the new ring road of Volos and Agria opened over a distance of two kilometers perfect outcrops through the Triassic-Jurassic marble formation (Fig. 12). Almost over the entire distance the marbles are fragmented, ragged and chattered. Monomict breccia are predominant, in the upper part meter-size blocks frequently are present. Several portions were highly fissured and filled with oxide- rich calcite cement, thus excluding the effect of brecciation by blasting. Further brecciation is present within the marble formation along the new roads from Agria to the village Drakia, from Kato Lechonia to the villages of Servanates (Fig. 13a-d) and Aghios Lavrendios, and from Ano Lechonia to Aghios Vlassios. In all cases, especially the surface of the marbles is highly fractured and exhibit strong striation. These features are repeated at the coastal cliffs of Rivera beach west of Kato Ghadzea (Fig. 13e&f). There, the monomict breccias are filled with reddish Pleistocene red beds and hydrothermal deposits, dipping towards sea. The gently undulating peneplain morphology with altitudes between 300 and 400m of south-central part of the Pelio mountain chain differs significantly from that of the mountainous, up to 1500m high northwestern part, cut by steep valleys, probably due to strong differential tectonic uplift during Tertiary times. The southern part is made up of thrust sheets of uncertain Triassic to Jurassic age, which are composed of 8 epidote-actinolite-chlorite schists, mica schists, phyllites, crystalline limestones and platy marbles (Figs. 6,7), as well as major metamorphosed mafic bodies and small lenses of highly tectonized serpentinites. Characteristic outcrops in the area between the village of Afetes and the plain of Myriovrity, which show the uppermost Quaternary surface in Fig. 14, demonstrate the style of unorientated deformation and chaotic brecciation by shearing and late faulting overprinting early thrusting during orogenic phases and uplift. Similar aspects occur in the same tectonic units along the gulf coast between Afissos and Milina. Characteristic monomict and less polymict breccia are present north and south of Lefokastro and Paou beach (southwest of Arghalasti, Fig. 9). The surface layers of crystalline limestones are chaotically brecciated (Figs. 24,25) and cut by irregular breccia channels cemented with matrix calcite and oxides. In contrast, schists are and gneissic layers are only fractured; serpentinite lenses are not affected. Cataclastic fragmentation and brecciation are characteristic features below the present surface in the Triassic to Jurassic marble formation in the Tiseo Mountains. The coast between Lefokastro and Paou and the new road cuts along the gulf coast between Chorto and Trikeri bay show some impressive examples (Figs. 15,16). Matrix supported monomict breccias are abundant between Marathias bay and Avra. A few hundred meters west of the chapel Panaghia a local trough in the mountain side wall next to the road leads to an impression of a local impact phenomenon, since the white marble layers are highly compressed and pulverized, the trough itself filled with brecciated red bed material (Figs. 176a-d). A geomorphological strange almost circular feature is present in the 600x700m wide plain of Xerokabos at 14 meters above sea level between Marathias in the west and Chondhri Amos Bay in the East at the Aegean Sea. This plain was interpreted as a large doline in the Geological map of Greece (Sheet Argalasti, Katsikatsos et al. 1989). However, searching for evidence of a doline or an impact crater turned out to be negative. The style of fragmentation and brecciation changes due to the lithologies of the opposite mountains of Stefanania and Dhiasela in the Trikeri peninsula. There, the white marble formation of the Tiseo Mountains is overthrusted by a complex zone of thrust sheets consisting of cherts, shales, schists, crystalline limestones and ophiolitic relicts, which turns into a thick sequence of Upper Cetaceous platy marbles and cherts (Figs. 18). In many places along the road cuts up to the village of Trikeri (Figs. 19b-d), the wide overthrust zone is extremely fractioned and chattered including all grain sizes and large blocks in a loose unconsolidated sandy matrix, covered by very thin layers of soil, bushes and trees. These morphological features of cataclastic fragmentation and brecciation continue to an extreme in the island of Palea Trikeri (Figs. 19a-d) and in the islets of

9 Strongyli, Psathi and Pithou, the latter investigated from the sea (Figs. 19e&f). The upper few meters below the soil and vegetation seem to be shattered with fragmentation down to centimeter dimensions. All lithologies are dipping towards the center of the gulf.

Figs. 10 Outcrops on the northwestern coast between Nea Anchialos, Cape Agistri and the hill of Amphanae (Gulf of Volos). (a) Roadcut 3 km east of Nea Anchialos. Greenschists below surface heavily brecciated; (b) between Kritharia and Stavros distorted and brecciated crystalline limestones and schists, partly filled with reddish Pleistocene red beds and hydrothermal deposits, dipping towards sea; (c) and (d) Cape Agistri green broken gneisses with crystalline limestone 10 intercalations; € and (f) Fractured gneiss quarry below hill of Amphanae with heavily broken surface. Photos V. Dietrich.

Figs. 11 Shore along Cape Ghoritsa. (a)-(c) Roadcuts showing different types of fracturing the Triassic marbles from simple fracturing to cataclastic monomict breccia and in part pulverization. Part of it is seen as a result of blasting during construction. However, outcrops at the seaside show the same effects (d). In all cases, the surface below the soil show consistently fine to moderate monomict brecciation. Outcrop (e) at sea level exhibit a network of calcite and quartz veins as well as an open conjugate network of straight fractures (f). Photos V. Dietrich.

Figs. 12 Highway construction 2009 above Asteria (Agria, northeast coast Pagasitic Gulf). The road cuts the Upper Triassic to Jurassic marbles of the lower Pelagonian nappes in the Kastri Mt. 429m. a-d show the brecciation at all scales with meter size blocs (c and d), which may be in parts due to tectonic emplacement and uplift during the orogenic phases at shallow crustal level and as well as due to blasting besides a possible impact. The latter seems to be more evident in (e and f), indicating high fragmentation and healing with oxide-rich calcite cement. Photos V. Dietrich.

Figs. 13a-d Roadcuts below the village of Servanates (above Kato Lechonia); e and f at Rivera beach west of Kato Ghadzea. In all Figures strong surface deformation in marble intercalations of the Upper Paleozoic to Triassic thrust sheets; c and d show strong internal brecciation, initiated by a hydrothermal overpressure event; e and f exhibit cataclastic brecciation and filled with reddish Pleistocene red beds and hydrothermal deposits, dipping towards sea. Photos V. Dietrich.

Figs. 14 Roadcuts between the town of Arghalasti and Miriovriti through the Quaternary surface of the complex thrust sheets of Upper Jurassic schists, greenschists and phyllites. Characteristic features are non-oriented deformation effected by shearing and late faulting in addition to early thrusting during orogenic phases and uplift (a-d, north of Afetes); (e) road below Alkopetra Point 978m, 8.4km SW Chamia; (f) Road next to shore, east coast of Kokkinovrachos; show lenses of highly sheared light greenish serpentinite. The surface is covered by up to one meter of soil. Photos V. Dietrich.

Figs. 15 North and South of Paou beach (southwest of Arghalasti). The coastal outcrops of monomict breccia are representative for many of their kind along the western shores between the villages of Afissos and Milina. They are most significant in platy crystalline limestones. The surface layers are chaotically brecciated (a and b) and cut by irregular breccia channels cemented with matrix calcite and oxides (c-f). In contrast, schists are and gneissic layers are only fractured; serpentinite lenses are not affected. Photos V. Dietrich.

Figs. 16 Coastal outcrops at Milina, Razi and Lefokastro. Chaotic, monomict, clast supported crystalline limestone breccia (a and b) with blocs up to meter size occur of the entrance of the village of Milina. (c-e) monomict and polymict matrix supported breccia containing schist fragments at Razi; in (d) the breccia is topping broken layers of platy limestone. (f) is a monomict matrix supported breccia at Lefokastro. Photos V. Dietrich.

Figs. 17a-d Roadcuts west of Panaghia chapel (Trikeri Bay) with a unique deformational geomorphological feature. The brecciated Quaternary surface seems to fill highly deformed depression in pulverized Triassic marbles (detail in d). This effect cannot be explained by any recent dynamite blasting, because of the undisturbed dense vegetation above the outcrops up the hill. The only interpretation remains at the moment due to an impact effect; (e and f) show polymict, matrix supported breccia cemented with calcite and oxides along the road along the shore between Avra and Marathias south of Milina. Photos V. Dietrich.

Figs. 18 Roadcuts above Kotes northwest of Trikeri Village. Unconsolidated chaotic matrix supported breccia, in parts polygenic with crystalline limestone, metasandstones, and cherts fragments (a-d). These occurrences are present in the entire eastern slopes and contain larger lenses of sheared serpentinites; (e) represent the uppermost brecciated surface of the Upper Cretaceous platy crystalline limestones on the western slopes of Trikeri Peninsula, west of Trikeri Village; (f) a brecciated part on the road down to the northern ferry station to Old Trikeri Island. Photos V. Dietrich.

Figs. 19 Breccias at Old Trikeri Island and Pitou Islet; (a) View of Old Trikeri Island and Pithou islet from north. In the northern and western parts of the island (e.g. Cape Pardhalos; (b) thrust sheets consisting of red-brownish colored alternations of mica schists, phyllites, metasandstones, cherts, (e.g. c at Prasini Amos), dark-colored crystalline limestones with greenish lenses of highly tectonized serpentinites and meta pillowlava (d); (e and f) same lithologic situation at Pitou Islet; large green serpentinite lenses embedded in reddish schists and cherts with minor overthrust horizons. All lithologies in low-greenschist facies are deeply sheared, fractured, mylonitized. Photos V. Dietrich.

2.1.1 Striation and Shatter Cones

Marbles and crystalline limestones of Triassic to Upper Cretaceous age encompass a large portion of the lithologies of the surroundings of Pagasitic Gulf, which inhibit multiple generations of deformation, effects of nappe emplacement, thrusting and uplift in the crust during Tertiary times. In many localities monomict breccias occur as results of very young deformation and fracturing events, probably Neogene or Holocene of age (e.g. Fig. 20a). Beside the breccia, many lithologies must have been exposed at the surface during Pleistocene times. Metamorphosed calcareous rocks, marbles as well as platy crystalline limestone show characteristic striation features on more less plain surfaces, which resemble divergent sets of micro-channels as erosional results of heavy rain. Surficial striation is generally known as long scratches and gouges on polished rock surfaces by glacial abrasion. However, on fresh surfaces and major broken blocks they turn out to be an erosional surficial effect of an internal striation (Figs. 20b-f). Penetrative straight and curved, closed or wide spaced surficial striations in metamorphic rocks are mainly a result of intersections of sets of healed planar fractures, in the case of metamorphic and tectonized limestones and marbles dislocations of twin-lamellae (Figs. 20). Here, irregular, slightly curved striation penetrates the crystalline limestone as well as marble, and thus are regarded first order as metamorphic effects, similar but not identical to typical “shatter cones” found in bedrocks and ejecta in several impact craters as solid evidence for shock metamorphism (French and Short 1968; French 1998; French and Koerbel 2010; Langenhorst 2002; Buchner 2018). Shatter cones are described in detail as “a range of curved to curvilinear fractures decorated with more or less divergent striations. Striations radiate from an apex of a conical feature or from a narrow (a few millimeters to a few centimeters wide) apical area” (Nicolayson and Reimolds 1999). Shatter cones may have formed by the interference of shock waves at shock pressures above 2 GPa (Baratoux and Melosh 2003) and therefore appear to be a macroscopic shock indicator representative indication of low-grade shock metamorphism.

Figs. 20 Surficial deformational pattern of quartz lenses, crystalline limestone and marble layers. (a) Quartz block highly fractured including two sets of closed space parallel fractures; road at Paou monastery; (b and c) weathered surfaces of crystalline platy limestones with narrow divergent ridges (striation) as result of deformation; localities Prasini Amos, Old Trikeri Island and close to the port; (d) divergent fractures emanating from small apex areas in a fresh limestone surface at Rivera beach and at the shore line of Cape Ghoritsa (e). They resemble curvilinear fractures of “shatter cones” and thus seen as a macroscopic shock indicator of low-grade shock metamorphism; (f) fresh broken white marble block with internal striations; locality road cut through the southern slope of Chteni Mt. above Agria. Photos V. Dietrich

21 2.2 Microscopic Scale 2.2.1 Tectonic and Planer Deformation in Quartz

Sampling was concentrated on quartz veins in all lithologies along the shores and hill sites around the Pagasitic Gulf, thus in the distal environment of the possible crater formed by an impact blast. The goal was to detect impact related shock metamorphic deformation. In order to discriminate post-orogenic shock-induced cataclastic fabrics, orogenic tectonic fabrics had to be recognized and extracted. In most cases, multi deformational phases are present overprinting each other, which obliterates a proper identification. Undulate extinction and “Boehm lamellae” (metamorphic deformation lamellae, MDLs) in quartz-fabrics as a progression of unorganized dislocations are the most common deformational fabrics and are regarded as results of tectonic deformation in the crust during orogenic processes (Böhm 1883; Christi and Raleigh 1959; Christie and Ardell 1974). Examples of metamorphic fabrics in quartz-rich schists, gneisses and veins are shown in Figs. 21 and 22. In contrast, planar deformation fractures (PDF’s) in quartz and feldspars with a general thickness <1m and amorphous composition of the host mineral are the most diagnostic shock-indicators for high-pressures starting at 10-15 GPa. They are also recognized in other silicate minerals such as olivine and zircon. According to shock- experiments, PDF’s transform into diaplectic glass at 35 GPa. The speed of the shock event inhibits the timing for solid-state transformation of quartz and feldspars to their high-pressure polymorphs stishovite and lingunite (French and Short 1968; Stöffler and Langenhorst 1994; Langenhorst and Deutsch 1998, 2012; Langenhorst 2002; French and Koeberl 2010; Reimold and Jourdan 2012) The “decoration” is the result of the amorphous lamellae and precipitation of fluid bubbles (water). Clear sets of close spaced “planar deformation fractures“ (PDFs) have not been found up to date. Sub-planar and sub-parallel deformation lamellae, as well as weak “planar fractures (PFs)” with approx. 1m thickness, have been detected in some quartz crystals (Figs. 23 a-d), which are interpreted as a result of pressure shock-induced metamorphic effects. In addition, cases of planer fractures decorated with tiny unidentified inclusions may represent the effects of contemporaneous fluid circulation. Similar PFs have been described from several meteorite craters.

Figs. 21 Examples of quartz-fabrics in brecciated metamorphic limestone and in quartz veins caused by tectonic and metamorphic deformation. All quartz crystals show undulose extinction as a progression of unorganized dislocations. (a) brecciated sample mylonitised crystalline limestone with highly deformed quartz components (indicated with red arrows) crosscut by a broken calcite vein; road north of Trikeri village at Piridhistra; (b) Highly deformed large quartz grain with chaotic undulose extinction surrounded by moertel quartz grains; west of Mivrioti; (c and d) Flattened and elongate quartz grains healed by new horizons of intergranular micro quartz grains due to dislocation and migration; Kalamos beach west of Arghalasti; (e) Sets of metamorphic deformation lamellae (MDL’s): 1/2 fractured quartz, 3 typical closed spaced “Boehm lamellae”, 4 late open-spaced and healed fracture; road north of Xerokastros plain (Trikeri); (f) sets incomplete tensional planar fractures overprinting earlier deformed quartz with MDL’s. All images under crossed polarized light.

Figs. 22 Deformational fabrics in quartz-plagioclase veins. (a and b) close spaced wavy fractures with offset dislocations overprinting undulose extinction; (c) plagioclase grains with incomplete close parallel planar fractures overprinting original twinning; (d) wide spaced conjugate planar fractures in plagioclase with inclined offset; both samples at coastal road east of Nea Anchialos. Images (a-c) under crossed polarized light, (b) under plain polarized light.

Fig. 23 Close spaced planar fractures (PF’s) in quartz veins; (a) close spaced planar fractures next to quartz grain with ballen-type deformational fractures (left upper part of image; from the upper schists of Neochori; (b) close spaced planar fractures overprinting prior deformed quartz-fabrics with curved undulose extinction; east coast road of Kokkinovrachos; (c and d) quartz-fabrics with straight close spaced planar fractures with approx. 1m thickness in different quartz grains in same sample at coastal rod east of Nea Anchialos. All Images under crossed polarized light.

2.2.2 Tectonic and Planer Deformation in Calcite

Apparently, calcite deformational effects can easily be produced in target limestones by impact shock-induced pressures (Baratoux and Melosh 2003; Burt et al. 2005, in the Arizona Barringer Crater). Multiple twinning in calcite is a common feature and originates from metamorphic deformation and recrystallization of limestones to marbles at low pressures and temperatures. It also occurs in coarse calcite crystals as products of hydrothermal or metasomatic fluids in the upper crust and near surface horizons, respectively. In addition, calcite shows a predominant cleavage usually in three directions parallel to the rhombohedron. Calcite cleavages as well as multiple twins are easily deformed by dynamic tectonic processes, such as by shearing and folding in faults, thrusts and folds (for comparison, see Figs. 24a-d).

25 Therefore, low- and medium-density twins in calcite formed by shock metamorphism, cannot be reliable distinguished from those formed by tectonic deformation; only high-density twins and planar fractures (PFs) that offset cleavage planes, may be regarded as impact related-shock effects (Langenhorst et al. 2000; Burt et al. 2005; Huson et al. 2009, 2011; Hamers and Drury 2011). Consequently, these deformational effects can be recognized in the XRD powder pattern of calcite and dolomite in “Single Peak Profiling - Full Width Half Maximum” widening (Huson et al. 2009, 2011). Planar fractures (PFs) have been found in several cases in the crystalline limestones of low metamorphic grade as well as in the marble of the Triassic-Jurassic formation (Figs. 24e&f). The very closed spaced parallel fractures have a thickness <1m and spaces between them of 1-5m, whereas wide spaced straight parallel planar fractures are 5-20 m thick. These cataclastic, post-Alpine calcite deformation fabrics are compatible to the observed planar fractures in quartz grains and thus, indicative for an impact-induced event. Macroscopic aspects of striation as result of intersections of sets of healed planar fractures have been discussed in the previous chapter 2.2.2.1 and their resemblance to the curvilinear fractures of shatter cones.

Figs. 24 Chaotic deformation fabrics in calcite. (a) Large deformed calcite grain rimmed by isometric micro calcite moertel grains; sharp planar close spaced fractures (red marker) overprinting typical calcite twinning; component in breccia Razi north of Lefokatro; (b) Large deformed calcite grain with sharp planar close spaced fractures (red marker) next to quartz grains with “Boehm lamellae”; road down from Portaria above Volos. (c) highly deformed multiple twinned calcite elongation and quartz (yellow colors) in marble at beach cliffs Rivera; (d) close spaced planar calcite winning with amoeboid quartz intergrowth (sea cliff Cape Ghoritsa); (e) highly deformed marble with multiple sets of close spaced lamellae and fractures with red markers: 1) slight bend lamellae, 2) very closed spaced parallel fractures <1m and spaces between 1-5m; 3) wide spaced straight parallel planar fractures (construction highway above Asteria/Agria); (f ) twinned calcite crystals No 1, cut by sharp planar close (No. 2) and wide spaced fracture planes with a thickness between 5-20 m (No. 3) and wide spaced fracture planes (No. 3) red marker); road cut at monastery Paou. e) image under plane polarized light, all other images under crossed polarized light.

27 2.3 Metamorphic and Post-Alpine Hydrothermal Activity (Veining)

Fluid circulation must have occurred during and subsequent to any impact or blast within the crater fillings and in the surrounding area due to intensive fractionation and brecciation. The interaction of shock-melted or heated target rocks with surficial meteoric and/or aquifer underground water will induce hot-rock-water circulation, which, depending on the temperature gradient, can dissolve, transport and precipitate the equivalent hydrothermal mineral paragenesis. Post-impact hydrothermal alteration has been described from several impact craters (Osinski e al. 2001). In the case of the Tertiary large Ries impact crater, hydrothermal fluids were generated as a mixture of meteoric water with dissolved components from glassy suevites of the crater fill (Osinski 2005). Calcite and quartz precipitation led to the most common filling of veins in all lithologies in the surroundings of the Pagasitic Gulf. The occurrence of different generations shows that a large part of veins had been formed already during orogenic events originating from metamorphic fluids in the crust (Figs. 25a-d) prior to any impact of blast. However, post-tectonic veining with calcite/aragonite and quartz precipitation are also present, which is evident in large fractures and in open cavities (Figs. 25 e-f and 26 a-f). In addition, large open fracture systems combined with breccias are filled with surficial “red bed” material, probably Pleistocene in age that consists of mixtures of smectite, calcite and iron and manganese oxides (goethite and hematite).

Fig. 25 Quartz and calcite veins in schists, marbles and crystalline limestones as result of crustal metamorphic fluids under tensional stresses. (a) Green mica-schists with brownish thin deformed marble layers and white quartz lenses and horizons; beach of Lefokastro south of Afissos; (b) Different Quartz vein generations in fractured limestone along the east coast road of Kokkinovrachos; (c-e) Rivera beach cliffs. c Echellon set of calcite veins in fractured crystalline limestone; (d) Weathered surface of crystalline limestone with divergent striation cut by a brownish calcite vein. e Wide spaced fractures in crystalline limestone filled with matrix supported calcite breccia. f Wide spaced fractures in marble filled with matrix supported calcite breccia at Cape Goritsa. Photos V. Dietrich.

Fig. 26 Different types of late hydrothermal activity in the distal environment. a Wide spaced fractures in crystalline limestone filled with oxidized matrix supported calcite breccia (Cape Agistri); b Close up polymict matrix supported breccia with components of dark crystalline limestone and white marble (Servates above Lechonia); c Matrix supported polymict breccia at beach of Kritharia west of Cape Agistri; d Open cavity and fractures filled with aragonite and calcite hydrothermal deposits (Servates above Lechonia); e and f Hydrothermal aragonite deposits in open cavities along road cuts through the western slopes of Koutra (west of Trikeri Village). Photos V. Dietrich.

30 References

Baratoux, D., Melosh, H. J. (2003). The formation of shatter cones by shock wave interference during impacting. Earth and Planetary Science Letters 216, 43–54 Böhm, A. (1883). Ueber die Gesteine des Wechsels Mineralogische und petrograph. Mitthlg. von G. Tschermak. 5, 197-214 Burt, J.B., Pope, K.O., Watkinson, A.J. (2005). Petrographic, X- ray diffraction and electron spin resonance analysis of deformed calcite: Meteor Crater, Arizona. Meteoritics and Planetary Science 40, 296-305 Caputo, R. (1990). Geological, structural study of the recent, active brittle deformation of the Neogene-Quaternary basins of Thessaly (Greece), Scientific Annals, 12, Aristotle University of Thessaloniki, 2, 252 p., Thessaloniki Caputo, R., Pavlides, S. (1993). Late Cenozoic geodynamic evolution of Thessaly and surroundings (central-northern Greece). Tectonophysics 223, 339-362 Christie, J. M., Ardell, A. J. (1974). Substructures of deformation lamellae in Quartz. Geology, 2,8, 405-408 Christie, J. M., Raleigh, C. B., (1959). Origin of deformation lamellae in quartz. American Journal of Science 257, 385-407 Collins, G. S., Melosh, H. J., Marcus, R. A. (2005). Earth Impact Effects Program: A Web-based computer program for calculating the regional environmental consequences of a meteoroid impact on Earth. Meteoritics & Planetary Science 40, 6, 817-840 Craddock, J.P., Klein, T., Kowalczyk, G., Zulauf, G. (2009). Calcite twinning strains in Alpine orogen flysch: Implications for thrust-nappe mechanics, geodynamics of Crete. Lithosphere 1, 3, 174-191 Dietrich, V.J., Lagios, E., Reusser, E. Sakkas, V., Gartzos, E., Kyriakopoulos, K. (2017). The Zerelia Twin-Lakes (Central Greece) Two Possible Meteorite Craters. 116p., LAP LAMBERT Academic Publishing. ISBN-13: 978-3-659- 64693-5 Ferrière J. (1977). Le secteur méridional du “massif métamorphique de Thessalie”. Le massif du Pélion et ses environs. Proceedings of the VI Colloquium on the Geology of the Aegean region, 291-309 Ferrière, J. (1976). Etude préliminaire des terrains métamorphiques de la presqu’ile du Pelion anterieurs aux niveaux conglomératiques présumés Crétacé superieur (Grèce continentale orientale). Consequences tectoniques. C R Acad Sci Paris 282,1485-1488 Ferrière, J. (1982). Paleogeographies et tectoniques superposees dans les Hellenides Internes au niveau de l’Othrys et du Pelion (Grèce). Soc Géol Nord Publ 8,1- 970 Ferrière, J., Baumgartner, P.O., Chanier, F. (2017). The Maliac Ocean: the origin of the Tethyan Hellenic ophiolites. Int J Earth Sci (Geol Rundsch), DOI 10.1007/s00531-016-1303-6 French, B. M., Koeberl, C. (2010). The convincing identification of terrestrial

31 meteorite impact structures. What works, what doesn't, and why. Earth- Science Reviews 98, 123-170 French, B.M., Short, N.M. (Eds) (1968). Shock Metamorphism of Natural Materials. Baltimore, Maryland, USA: Mono Book, 644 pp Galanakis, D., Pavlides, S., Mountrakis, D. (1998). Recent brittle tectonics in Almyros - Pagasitikos, Maliakos, N. Euboea, Pilio. Bulletin of the Geological Society of Greece 30, 263-273 Hamers, M.F., Drury, M.R. (2011). Scanning electron microscope- cathodoluminescence (SEM-CL) imaging of planar deformation features and tectonic deformation lamellae in quartz. Meteoritics and Planetary Science 46, 12, 1814-183 Huson, S., Pope, M., Watkinson, A.J., Foit, F. (2011). Deformational features and impact-generated breccia from the Sierra Madera impact structure, west Texas. Geological Society of America Bulletin 123, 1-2, 371-383 Huson, S.A., Franklin, F., Foit, F.F., Watkinson, A.J., Michael, C., Pope, M.C. (2009). Rietveld analysis of X-ray powder diffraction patterns as a potential tool for the identification of impact-deformed carbonate rocks. Meteoritics and Planetary Science 44, 11, 1695-1706 Jacobshagen, V., Skala, W., Wallbrecher, E. (1979). Alpine structure, development of the southern Pelion Peninsula, the North Sporades. in: Mediterranean Orogens - Geodynamic Investigations along Geotraverses through Alps, Apennines, Helenides; Eds. H. Closs, D.H. Roeder, K. Schmidt; Schweizerbart, 484-488, Stuttgart 1978 Jacobshagen, V., Wallbrecher, E. (1984). Pre-Neogene nappe structure, metamorphism of the North Sporades, the southern Pelion peninsula. Special Publication of the Geological Society of London 17, 591-602 Katsikatsos, G., Migiros, G., Triantafphyllis, M., Mettos, A. (1986). Geological structure of Internal Hellenides (E. Thessaly-SW Macedonia-Euboea-Attica- Northern Cyclades Islands, Lesvos). IGME, Geol. Geoph. Res., Special Issue, 191-212 Katsikatsos, G., Mylonakis, I., Triantaphylus, E., Papadeas, G., (1989). IGME 1:50000 Geological Map of Greece, Sheet Argalasti, IGME Athens Katsikatsos, G., Mylonakis, I., Vidakis, M., Hecht, J., Papadeas G., (1978). IGME 1:50000 Geological Map of Greece, Sheet Volos, IGME Athens Katsikatsos, G., Papadeas, G., Mylonakis, I., Triantaphylus, E. (1987). 1:50000 Geological Map of Greece, Sheet Zagora - Syki, IGME Athens Kenkmann, Th., Poelchau, M.H., Wulf, G. (2014). Structural geology of impact craters. J. Structural Geology 62, 156-182 Korres, G., Triantafyllou, G., Petihakis, G., Raitsos, D.E., Hoteit, I., Pollani, A., Colella, S., Tsiaras, K. (2011). A data assimilation tool for the Pagasitikos Gulf ecosystem dynamics: Methods, benefits. Journal of Marine Systems. Doi: 10.1016/j.jmarsys.2011.11.004 Kreemer, C., Chamot-Rooke, N., Le Pichon, X. (2004). Constraints on the evolution, vertical coherency of deformation in the Northern Aegean from a comparison

32 of geodetic, geologic, seismologic data. Earth Planetary Science Letters 225, 329-346 Langenhorst, F. (2002). Shock metamorphism of some minerals. Basic introduction and microstructural observations. Bulletin of the Czech Geological Survey 77, 4, 265-282 Langenhorst, F., Deutsch, A. (1998). Minerals in terrestrial impact structures and their characteristic features. In Marfunin A.S. (Ed.), Advanced Mineralogy 3, Springer, Berlin, 95-119 Langenhorst, F., Deutsch, A. (2012). Shock metamorphism of minerals. Elements, 8, 1, 31-36 Langenhorst, F., Deutsch, A., Ivanov, B.A., Hornemann, U. (2000). On the shock behaviour of CaCO3. Dynamic loading and fast unloading experiments - modelling - mineralogical observations. Lunar and Planetary Science Conference 31, abstracts, 1851 Lips, W., Wijbrans, J. R., White, S. H. (1999). New insights from 40 Ar/ 39 Ar laserprobe dating of white mica fabrics from the Pelion Massif, Pelagonian Zone, internal Hellenides, Greece; implications for the timing of metamorphic episodes, tectonic events in the Aegean region. Special Publication of the Geological Society of London 156, 457-474 Marinos, G., Anastopoulos, J., Maratos, G., Melidonis, N., Andronopoulos, B. (1957). Geological maps of Greece, 1 : 50 000, sheet Anavra, Institute for Geology , Subsurface Research, Athen. Marinos, G., Anastopoulos, J., Maratos, G., Melidonis, N., Andronopoulos, B. (1962). Geological Map of Greece 1: 50 000, Sheet Almiros. Institute for Geology, Subsurface Research, Athen. Nance, R.D. (2010). Neogene-Recent extension on the eastern flank of Mount Olympos, Greece. Tectonophysics 488(1-4), 282-292 Nicolaysen, L. O., Reimold, W. U. (1999). Vredefort shatter cones revisited. Journal of Geophysical Research 104, 4911-4930 Osinski, G.R, Spray, J.G, Lee, P. (2001). Impact-induced hydrothermal activity within the Haughton impact structure, Arctic Canada; generation of a transient, warm, wet oasis. Meteoritics & Planetary Science, 36, 731-745. Osinski, G.R. (2005). Hydrothermal activity associated with the Ries impact event, Germany. Geofluids 5, 202-220 Papanikolaou, D., Alexandri, M., Nomikou, P. (2006). Active faulting in the North Aegean basin. In Dilek, Y., Pavlides, S. (Eds.). Postcollisional tectonics, magmatism in the Mediterranean region, Asia: Geological Society of America Special Paper 409, 189-209 Papanikolaou, I., Migiros G. (2008). Brittle Deformation, Hydrogeological Pattern of The Eastern Pelion Area (Tsangarada). 8th International Hydrogeological Congress of Greece - 3rd MEM Workshop on Fissured Rocks Hydrology. (Eds. G. Migiros, G. Stamatis, G. Stournaras) Athens 2008, 1, 347-360

33 Papanikolaou, I.D., Papanikolaou, D.I. (2007). Seismic hazard scenarios from the longest geologically constrained active fault of the Aegean. Quaternary International 171-172, 31-44 Papazachos, C. B. (1998). Crustal P- and S-velocity structure of the Serbomacedonian Massif (Northern Greece) obtained by non-linear inversion of traveltimes, Geophys. J. Int., 134(1), 25–39 Papazachos, B. C, Papazachou, C. B. (1989). The earthquakes of Greece. Ziti Publications, Thessaloniki, 356 pp. Perissoratis, C., Angelopoulos, I., Mitropoulos, D., Michailidis, S. (1991). Surficial sediment map of the Aegean Sea Floor: Pagasitikos Sheet, scale 1 : 200’000, ed. IGME, Athens. Petihakis, G., Tsiaras, K., Triantafdyllou, G., Korres, G., Tsagaraki, T.M., Tsapakis, M., Vavillis, P., Pollani, A., Frangoulis, C. (2012). Application of a complex ecosystem model to evaluate effects of finfish culture in Pagasitikos Gulf, Greece. Journal of Marine Systems 94, supplement, S65-S67 Reimold, W.U., Jourdan, F. (2012). Impact! – Bolides, craters, and catastrophes. Elements 8,1, 19-24 Sachpazi, M., Galve, A., Laigle, M., Him, A., Sokos, E., Serpetsidaki, A., Marthelot, J.M., Alperin, J.M.P., Zelt, B., Taylor, B. (2007). Moho topography under central Greece and its compensation by Pn time-terms for the accurate location of hypocenters: The example of the Gulf of Corinth 1995 Aigion earthquake, Tectonophysics, 440 (1-4), 53–65 Sakellariou, D., Galanidou, N. (2015). Pleistocene submerged landscapes, Palaeolithic archaeology in the tectonically active Aegean region. Geological Society, London, Special Publications 411,1, 145-178 Sakellariou, D., Rousakis, G., Kaberi, H., Kapsimalis, V., Georgiou, P., Kanellopoulos, TH., Lykousis, V. (2007). Tectono-sedimentary structure, late quaternary evolution of the north Evia gulf basin, Central Greece: preliminary results. Bulletin of the Geological Society of Greece 37, 451-462 Schenker, F.L., Burg, J.-P., Kostopoulos, D, Moulas, E., Larionov, A., von Quadt, A. (2014). From Mesoproterozoic magmatism to collisional Cretaceous anatexis: Tectonomagmatic history of the Pelagonian Zone, Greece, Tectonics, 33, 1552–1576, doi:10.1002/2014TC003563 Schenker, F.L., Fellin, M.G., Burg, J.-P. (2015). Polyphase evolution of Pelagonia (northern Greece) revealed by geological and fission-track data Solid Earth, 6, 285–302, doi:10.5194/se-6-285-2015 Schermer, E. (1993). Geometry, kinematics of continental basement deformation during the Alpine orogeny, Mt. Olympos region, Greece. J. Struct. Geol. 15, 571-591 Schermer, E.R. (1990). Mechanisms of blueschist creation, preservation in an A-type subduction zone, Mount Olympos region, Greece. Geology 18, 1130-1133 Schermer, E.R., Lux, D.R., Burchfiel, B.C. (1990). Temperature–time history of subducted crust, Mount Olympus region, Greece. Tectonics 9, 1165-1195 Stiros, S., Pirazzoli, P., Pomoni-Papaioannou, F., Laborel, J., Laborel, F., Arnold, M. (1994). Late Quaternary uplift of the Olympus - Pelion range coasts

34 (Macedonia - Thessaly, Greece). Bulletin of the Geological Society of Greece 30, 325-330 Stöffler, D., Langenhorst, F. (1994). Shock metamorphism of quartz in nature and experiment: I. Basic observation and theory. Meteoritics 29, 155-181 van Andel, T.H., Perissoratis, C. (2006). Late Quaternary depositional history of the North Evvoikos Gulf, Aegean Sea, Greece, Mar. Geol., 232,157-172 Wallbrecher, E. (1976). Geologie und Tektonik auf dem Südteil der Magnesischen Halbinsel (Nordgriechenland). Zt. dt. Geol. Ges. 127, 365-371, Hannover Wallbrecher, E. (1977). Nappe units of the southern Pelion Peninsula, their Origin. Proceed. VI Colloqu. Geol. Aegean Region, 281-290, Athen Wallbrecher, E. (1983). Alpidischer Deckenbau und Metamorphose auf den Nord- Sporaden und auf der Südlichen Pelion-Halbinsel, (Thessalien, Griechenland). Berliner Geowiss. Abh. (A) 48, 99-116, Berlin 1983

3 Geophysical Investigations of Pagasitic Gulf and Surrounding Areas Introduction

Applied Geophysics constitutes one of the basic tools in investigating terrestrial impact craters. The application of various geophysical techniques (gravity, magnetics, resistivity, seismics - especially reflection studies) may provide useful evidence about the interior structure of the impact craters, and furthermore to establish certain criteria for evaluating hypotheses of impact craters. However, those geophysical techniques can provide only supporting evidence; confirmation of the impact origin of such hypotheses is based on evidence produced by various geological laboratory disciplines. The most common and characteristic ground signature of an impact body is (usually) a circular gravity low corresponding to a bowl-shaped simple crater, which is interpreted as due to the presence of (allochthonous) breccia lenses in its lower part, or a more complex crater due to fractured parautochthonous rocks in the floor of the crater that did not escape the cavity by ejection. The size of these gravity anomalies increases with crater diameter (D), and the gravity anomalies may have amplitudes of -20 to -30 mGal, corresponding to crater diameters of 20 to 30 km (Pilkington and Grieve 1992). It appears that as crater diameter increases (D>30 km), the maximum negative gravity anomaly is reaching a limiting value of about -30 mGal, not having any further significant effect in the amplitude of the anomaly to be subsequently modelled (Pilkington and Grieve, 1992). The formation of the initial crater after an impact, called transient crater, is only an intermediate step in the development of the final crater size. The effect of the transient crater’s collapse should inevitably be considered for the formation of the final crater diameter. Craters smaller than 3 km in diameter are called (Dence 1965) simple craters; these are formed by highly brecciated and molten rocks at the base of the crater (Grieve et al. 1977). When considering transient crater diameter (D) with respect to its initial Depth extent, the following equation seems to be valid: D = 3.3xDepth (1)

However, the larger complex craters (see detailed accounts in eg. Pilkington and Grieve (1992)) seem to have shallower final Depth/D ratios compared to simple craters. According to Collins et al. (2005), the average depth of a complex crater is estimated by adopting the relationship of Herrick et al. (1997), based on the similarity in surface gravity between Earth and Venus, as well as to the fact that the Venusian crater data are subjected to less erosional effects than the Earthly crater data: Depth = 0.4D0.3 (2)

Nevertheless, when considering also the Lunar crater data, starting with the scaling law for the moon and taking into consideration the ratio of the surface gravity of Moon to Earth, the new equation predicts an even more shallower depth:

Depth = 0.294D0.301 (3)

Gravity Measurements and Modelling

The Gravity Anomaly Map of Greece (Lagios et al. 1995) was primarily considered and used in the present impact hypothesis for the broader area o Pagasitikos Gulf. The gravity data of the enriched Gravity Data Bank of Greece (Lagios et al. 1996) were used and reduced to gravity anomaly values. However, since there were no marine gravityobservations inside the Pagasitic Gulf, data from EGM08 (Pavlis et al. 2008) were extracted and incorporated for the compilation of the Gravity Anomaly Map (Fig. 27) after they were properly homogenized in terms of datum and terrain corrections. Therefore, all gravity values of the gravity map are based on the Gravity Reference Formula of 1967 (GRF’67) and referred to IGSN’71 (Hipkin et al. 1988). It can be seen that high gravity values (>50 mGal) are associated with the eastern part of Pagasitic Gulf, where the high-density metamorphosed formations prevail. However, going westwards and at the extreme western part of Pagasitic Gulf, the amplitude of the gravity values drop down to 10 mGal associated with the sedimentary sequence of the Thessaly Plain, while a 20 mGal low seems to dominate the gulf. The Residual Gravity Anomaly Map (Fig. 28) was formed for the modelling procedure. The regional gravity anomaly map of the broader area (Lagios et al. 1995; 1996), together with the eastwardly increased linear regional trend appeared in the reduced gravity anomaly map (Fig. 27) was initially taken into consideration. However, a filtering procedure that was also attempted here was proved more efficient for the compilation of the residual gravity anomaly map; a wavelength (> 30-35 km) of gravity anomalies corresponding to the size of the Pagasitic Gulf was subtracted here. It can therefore be seen that total amplitude of about -16 mGal characterizes the gravity anomaly closure inside the gulf. The implementation of the modelling procedure to further investigate the size of the likely transient crater associated with the probable impact hypothesis requires an estimated value of a density contrast between the bedrock (metamorphosed rocks approx. 2700-2720 Kg/m3) and the superficial deposits (about 2500-2600 Kg/m3). A value of 170 Kg/m3 density contrast was here considered, being generally consistent with other similar cases (Pilkington and Grieve 1992).

Fig. 27. Bouguer Gravity Anomaly Map of the broader area of Pagasitic Gulf. Contour interval 2 mGal.

Fig. 28. Residual Gravity Anomaly Map of the broader area of Pagasitic Gulf. Contour interval 2 mGal.

Fig. 29. Calculated Gravity Anomaly Model Map produced by a single interface of 170 Kgr/m3 density contrast. Contour interval 2 mGal.

The gravity effect of such model by assigning the aforementioned value of gravity contrast is shown in figure 30, together with the almost E-W profile crossing the gulf, along which the variation of the modelled interface can be seen. A fairly good match has been achieved between the calculated (theoretical) and observed (residual) gravity anomalies. The modelled U-shaped feature at the center of Pagasitic Gulf (Fig. 30) seems to have an extent of about 20 km, and a model depth extent of at least 5800 m. This diameter value has a discrepancy of only 3% according to the estimated value being resulted by equation (1). A more analytical 3-D model of the structure inside the gulf may be seen in figure 31.

38 Fig. 30. Variation of the gravity model interface indicating the depth and the areal extent of the assumed transient crater deduced by a density contrast of 170 Kg/m3 along the almost E-W profile shown in Fig. 29.

Fig. 31. A 3-D model of the assumed transient crater in the Pagasitic Gulf resulted by gravity modelling of the Residual Gravity Anomaly map assigning a single interface of 170 Kg/m3 density contrast.

On the assumption of the impact hypothesis, Pilkington and Grieve (1992) have considered theoretical curves between amplitude of gravity anomalies (Δg) produced by various impacts and their associated crater dimensions. Considering here the “model crater dimension” of 20 km (Fig. 30) - applying a density contrast of 170 Kg/m3), a value of about 17 mGal is resulting from the empirical curve (Δg vs Crater Model) for impact on crystalline rock formations. This value is certainly consistent with the 16 mGal observed amplitude value in figure 30 (discrepancy of only about 6%). Applying also the Impact Earth Program (Collins et al. 2005) developed (2017) by ITAP for Purdue University (USA), the following impact parameter-effects are resulted when considering a body (density 3000 Kg/m3) of 2 km diameter impacted on a crystalline basement at 50o impact angle with a typical speed of 17 km/sec: Transient Crater Diameter: 18.1 km Transient Crater Depth: 6.38 km Final Crater Diameter: 26.5 km Final Crater Depth: 793 meters The crater formed is a complex crater.

These values of the transient crater diameter and crater depth constitute a discrepancy of 10% and 11%, respectively, from our model parameters that are within acceptable limits, even though a

39 single interface was used in the gravity modelling. Also, the final crater diameter and depth (present day) are realistic when considering the size of the Pagasitic Gulf, where the maximum depth of sea water is about 100 m. Considering equations (2) and (3) on the Venusian and Lunar estimated values of crater depth, it seems that the Lunar value is closer (9%) to the Earth model value (5800 m) than the Venusian value (19%). Conclusion: The size, shape and the form of Pagasitic Gulf seems to be consistent to an impact case hypothesis, according to the existing geophysical data, however, not constituting a proof of this hypothesis. It should be pointed out that the density of the impact body applied above constitutes almost the upper limit of density assigned for asteroids (Wilson et al. (1999), which varies between 2270- 3100 Kg/m3, in contrast to meteorite density variation that is higher, 3480-3640 Kg/m3 (Chapman, 1995). Therefore, implementing again the above Impact Earth Program, this time for the lower limit of the density (2270 Kg/m3) for an assumed asteroid body of the same dimensions and orbital trajectory, the following complex crater dimensions are calculated: Transient Crater Diameter: 16.4 km Transient Crater Depth: 5.82 km Final Crater Diameter: 23.8 km Final Crater Depth: 768 meters

An alternative and more geologically realistic gravity model along the same profile (E-W) that would account for the depth and diameter of the final crater is shown in the following (Fig. 32), assuming the outlined densities of the conglomerates (2630 Kgr/m3) and the overlain soft sediments (2250 Kgr/m3). These values of crater dimensions shown above are realistically much closer to our estimated model values. That would imply that the Pagasitic Gulf seems to be consistent to an impact case hypothesis of an asteroid, rather than the case of a meteorite impact, the latter to be considered as rather less likely happening, not though constituting a proof of this hypothesis on the basis of the existing gravity data. However, from the size and shape alone, it is not possible to know if an impact crater was created by a comet or an asteroid.

The Air-burst Case Hypothesis

Because most comets are larger than 1km, they do create damage on the ground and many would cause impact craters. Jennisken and Popova (2018) provide a few accounts of major Earth’s craters formed by comets at different parts of the globe. Although the comet will be destroyed during its passage through the Earth’s atmosphere, it may be evaporated to form a gas-jet consisting of hot air and meteoric vapors. From this mixture which is decelerated over the length of the atmosphere, some gases can penetrate to the surface and cause the so-called surface airburst, which does not cause a crater. However, this hot jet-mixture can cause ground surface melting into glass layers of tens of centimeters thickness. Consequently, devastation and forest fires are inevitably expected to be taking place for a wider area (Shuvalov et al. 2016).

Fig. 32. An alternative more realistic gravity model of two interfaces along profile E-W (Fig. 33) for the impact case hypothesis.

Fig. 33. A 3-Dimentional gravity model of the top layer interface (Neogene) over Pagasitic Gulf based on the gravity values of Fig. 28 and the adopted densities shown in Fig. 32.

Fig. 34. A 3-D model of the assumed transient crater in the Pagasitic Gulf resulted by gravity modelling of the Residual Gravity Anomaly map of the bottom layer interface (corresponding to the transient crater interface).

At this point, though, it is necessary to point out that the entry angle of a body entering the Earth’s atmosphere is important and critical in this process. Not all the bodies entering the Earth’s atmosphere reach the ground surface. Depending on the entry angle and mass, they usually evaporate creating a vapor-jet. More specifically, a 100m-body in a vertical 50 km/sec impact, the formed vapor-jet will eventually reach the ground; however, at 45o and 30o entry angles, the created jets will be stopped at hights of 1-3 km and 6-9 km above the cellestrial surface, respectively. At a small angle of 5o, a 1000m-comet of the same entering speed will be totally evaporated, and the formed gas-jet reaches the ground at a much slower speed (Shuvalov and Trubetskaya 2007). An attempt was made to investigate whether the observed gravity anomaly over the wider area of Pagasitic Gulf (Fig. 28) could account for such a model of the upper crust, considering the characteristics of the top upper layers being suffered by a case of an air-burst. Such kind of model is presented along Profile E-W (Fig. 28) that was more extended on both sides over the gulf. Densities of 1900 Kg/m3 and 2500 Kg/m3 were assigned for the top two layers of yellow and red colour (Fig. 32), corresponding to Neogene sediments and the underlain sedimentary mixture consisting of thrust slices of Pelagonian napes, schists and limestones, respectively. The stippled third layer from the top, where a density of 2550 Kg/m3 was assigned, corresponds to the metamorphosed basement that was suffered by the hypothetical impact producing fishures and cracks in it. A density of 2700 Kg/m3 was finally assigned to the lower basement of the area that can be limestones or/and metamorphosed marbles. A borehole at the western part of the area on land, near Zerelia Lakes (Dietrich et al. 2017), was used as control of the sedimentary sequence in the above model, where the thickness of the Neogene sediments was larger than 350m, not reaching though the basement rock (Prof. K. Kyriakopoulos, personal communication). Inevitably, thus, larger thickness of the Neogenes of the top layer was resulted within the gulf (Fig. 32) compared to that one met at the borehole. 42

Fig. 32. A possible gravity model over an extendent profile on both sides of Pagasitic Gulf. Densities applied as in the text.

The outcome of the above model calculates a value of about 20km for the final crater diameter and a value of about 800m for the final crater depth. It therefore appears that the outcome of this model relating to the final depth (800m) and diameter crater (20km) dimensions is consistent with the results of the impact program outlined in the main text of the paper (717 m and 18.9 km, respectively), considering the air-burst hypothesis. The similar but shallower feature observed on the west of Pagasitic Gulf, which also results in the above modelling procedure, may rather be attributed to the probable/possible local basement depression.

Acknowledgements

Many thanks to Mr. S. Chailas for compiling the Gravity Anomaly Map of the broader area of Pagasitic Gulf.

43 References

Chapman, Clark R. (1995). «Galileo Observations of Gaspra, Ida, and Dactyl: Implications for Meteoritics». Meteoritics 30 (5): 496.

Collins, G. S., H. J. Melosh & R. A. Marcus, (2005). Earth impact effects program: A web-based computer program for calculating the regional environmental consequences of a meteoroid impact on Earth. Meteoritics & Planetary Science, 40(6), 817–840, doi:10.1111/j.1945- 5100.2005.tb00157.x.

Dence, M. R. (1965). The extraterrestrial origin of Canadian craters. Annual New York Academy of Science 123:941–969.

Dietrich, J.V., Lagios, E., Reusser, E., Sakkas, V., Gartsos, E. & Kyriakopoulos, K. (2017). The Zerelia Twin-Lakes (Central Greece): Two possible Meteorite Craters. LAP Lambert Academic Publishing. ISBN: 978-3-659-64693-5.

Grieve R. A. F., Dence M. R. & Robertson P. B. 1977. Cratering processes: As interpreted from the occurrence of impact melts. In Impact and explosion cratering, edited by Roddy D. J., Pepin R. O., and Merrill R. B. New York: Pergamon Press. pp 791–814.

Herrick, R. R., V. L. Sharpton, M. C. Malin, S. N. Lyons & K. Feely, 1997. Morphology and Morphometry of Impact Craters, in Venus II: Geology, Geophysics, Atmosphere, and Solar Wind Environment, edited by S. W. Bougher, D. M. Hunten, and R. J. Phillips, p. 1015.

Hipkin, R.G., Lagios, E., Lyness, D., & Jones, P. (1988). Reference gravity stations on the IGSN'71 standard in Britain and Greece. Geophys. J. Intern., 92, 143-148.

Jenniskens P.& Popova O. (2018). Comets in the path of the Earth. Element, 14, 107-112.

Lagios, E., Chailas, S. & Hipkin, R.G. (1995). Gravity and Isostatic Anomaly Maps of Greece produced. EOS, Transactions, American Geophysical Union, Vol. 76, 274 p.

Lagios, E. Chailas, S. & Hipkin, R.G. (1996). Newly compiled Gravity and Topographic Data Banks of Greece. Geophys. J. Intern., 126, 287-290.

Pavlis, N. K. , Holmes, S. A., Kenyon, S. C. & Factor, J. K. (2008). An Earth Gravitational Model to Degree 2160: EGM2008. EGU General Assembly 2008, Vienna, Austria, April 13-18.

Pilkington and Grieve, 1992. The Geophysical Signature of Terrestrial Impact Craters. Reviews of Geophysics, 30/2, 161-181.

Shuvalov V.V., Popova O.P., Svettsov V.V., Trubetskaya I.A. & Glazachev D.O. (2016). Determination of the height of the “Meteoric Explosion”. Solar System Research, 50, 1-12.

(Shuvalov V.V. & Trubetskaya I.A. (2007). Numerical modelling of impact induced aerial bursts. Solar System Research, 41/3, 220-230.

Wilson, L., Keil, K. & Love, S. J. (1999). The internal structures and densities of asteroids. Meteoritics & Planetary Science 34 (3): 479–483.