ΑΡΙΣΤΟΤΕΛΕΙΟ ΠΑΝΕΠΙΣΤΗΜΙΟ ΘΕΣΣΑΛΟΝΙΚΗΣ ΤΜΗΜΑ ΓΕΩΛΟΓΙΑΣ

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1 ΑΡΙΣΤΟΤΕΛΕΙΟ ΠΑΝΕΠΙΣΤΗΜΙΟ ΘΕΣΣΑΛΟΝΙΚΗΣ ΤΜΗΜΑ ΓΕΩΛΟΓΙΑΣ ΤΟΜΕΑΣ ΓΕΩΛΟΓΙΑΣ ΕΡΓΑΣΤΗΡΙΟ ΓΕΩΛΟΓΙΑΣ ΚΑΙ ΠΑΛΑΙΟΝΤΟΛΟΓΙΑΣ ARISTOTLE UNIVERSITY OF THESSALONIKI DEPARTMENT OF GEOLOGY ΓΕΩΜΕΤΡΙΑ ΤΗΣ ΠΑΡΑΜΟΡΦΩΣΗΣ ΚΑΙ ΚΙΝΗΜΑΤΙΚΗ ΑΝΑΛΥΣΗ ΣΤΗ ΜΕΣΟΕΛΛΗΝΙΚΗ ΑΥΛΑΚΑ GEOMETRY OF DEFORMATION AND KINEMATIC ANALYSIS IN MESOHELLENIC TROUGH Βαµβακά Αγνή (Vamvaka Agni) MSc Γεωλόγος ιδακτορική διατριβή (PhD Thesis) Thessaloniki 2009

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3 ΓΕΩΜΕΤΡΙΑ ΤΗΣ ΠΑΡΑΜΟΡΦΩΣΗΣ ΚΑΙ ΚΙΝΗΜΑΤΙΚΗ ΑΝΑΛΥΣΗ ΣΤΗ ΜΕΣΟΕΛΛΗΝΙΚΗ ΑΥΛΑΚΑ Υποβλήθηκε στο Τµήµα Γεωλογίας Τοµέας Γεωλογίας Ηµεροµηνία υποστήριξης: 5 Νοεµβρίου 2009 Επταµελής Εξεταστική Επιτροπή Κίλιας Αδαµάντιος Καθηγητής Τοµέα Γεωλογίας Α.Π.Θ. Μουντράκης ηµοσθένης - Καθηγητής Τοµέα Γεωλογίας Α.Π.Θ. Μιγκίρος Γεώργιος - Καθηγητής Τοµέα Γεωλογικών Επιστηµών και Ατµόσφαιρας Περιβάλλοντος του Γεωπονικού Πανεπιστηµίου Αθηνών Επταµελής Εξεταστική Επιτροπή Κίλιας Αδαµάντιος Καθηγητής Τοµέα Γεωλογίας Α.Π.Θ. Μουντράκης ηµοσθένης - Καθηγητής Τοµέα Γεωλογίας Α.Π.Θ. Μιγκίρος Γεώργιος - Καθηγητής Τοµέα Γεωλογικών Επιστηµών και Ατµόσφαιρας Περιβάλλοντος του Γεωπονικού Πανεπιστηµίου Αθηνών Σπύρος Παυλίδης Καθηγητής Τοµέα Γεωλογίας Α.Π.Θ. Ζεληλίδης Αβραάµ Καθηγητής Τµήµατος Γεωλογίας Πανεπιστηµίου Πατρών Ροντογιάννη-Τσιαµπάου Θεοδώρα Αν. Καθηγήτρια Τοµέα Γεωλογικών Επιστηµών του Μετσόβιου Πανεπιστηµίου Αθηνών Τρανός Μάρκος Λέκτορας Τοµέα Γεωλογίας Α. Π. Θ. Η έγκριση της ιδακτορικής ιατριβής από το Τµήµα Γεωλογίας της Σχολής Θετικών Επιστηµών του Αριστοτελείου Πανεπιστηµίου Θεσσαλονίκης δεν υποδηλώνει την αποδοχή των γνωµών του συγγραφέα (Ν. 5343/1932, 202 παρ.2)

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5 CONTENTS ABSTRACT ΠΕΡΙΛΗΨΗ PREFACE ΠΡΟΛΟΓΟΣ 1. INTRODUCTION Aim of the Study Methods used GEOLOGICAL SETTING Geomorphologic features General geological information Pindos Thrust Zone west boundary of the Mesohellenic trough Pelagonian microcontinent eastern boundary of MHT The Mesohellenic Trough (MHT) Proposed interpretations for the MHT evolution BASIN ANALYSIS BASIN TYPES CLASSIFICATION Basins Classification Subsidence Mechanisms FISSION TRACK THERMOCHRONOLOGY Principles of fission track thermochronology Age calculation Fission track annealing and modeling FISSION TRACK ANALYSIS IN MESOHELLENIC TROUGH AND THE ADJACENT PELAGONIAN BASEMENT Sampling and Analytical Procedure Thermochronological Data Results Pelagonian microcontinent samples (western part adjacent to MHT) Sediments samples of the Mesohellenic Trough... 40

6 5.4. Interpretation and Discussion Pelagonian microcontinent in the southern F.Y.R.O.M Pelagonian microcontinent in the study area (NW Greece) Sediments of the Mesohellenic Trough Conclusions on the FT analysis STRUCTURAL ANALYSIS Methodology Geometry and kinematics of deformation Field Observations Structures of Krania Formation at the Krania Spileo villages area (northern Krania sub-basin) Structures of Krania Formation at the Vassliki Nea Zoi villages area (southern Krania sub-basin) Structures of Eptachori and Pentalophos Formations area Structures of Tsotyli Formation area (eastern MHT margin) Tectonic events SUBSIDENCE RATES OF THE MHT STRUCTURAL EVOLUTION DISCUSSION BASIN TYPE CLASSIFICATION OF THE MHT REFERENCES APPENDIX A A.1. Fission Track Analysis Preparation of apatite samples A.2. Fission track counting and age determination APPENDIX B Stress inverse solutions of the Palaeostress analysis diagrams

7 ABSTRACT The Mesohellenic Trough (MHT) is a basin of more than 200 km length and km width, located in northern Greece and Albania. The MHT developed from Middle Eocene to Upper Miocene time, related to the Alpine orogenic processes, and is sited parallel to the structural fabric of the Hellenides (i.e. NNW-SSE), and between the Apulian microplate (External Hellenides, non-metamorphic) and the Pelagonian block (Internal Hellenides, metamorphic). The trough is described by its large size, great thickness of the sedimentary sequences (i.e., of 4km in single sections and much more in accumulative thickness), complicated structures, and numeral variations both in facies and thicknesses along and across the basin axis, comprising fan-delta conglomerates, alluvial and submarine fans (turbiditic sandstones and shales), deltaic and flood-plain sandstones and siltstones, and sandy shelf sediments. All these establish the Mesohellenic Trough as the largest and the most important late orogenic ( molasse-type ) basin in the Hellenides. Most of the previous works on MHT was focused on mapping (e.g., Brunn, 1956, 1960; Savoyat and Lalechos, 1969, 1972; Mavridis and Matarangas, 1979), sediment and facies analysis (e.g., Desprairies, 1979; Papanikolaou et al. 1977, 1988; Wilson, 1993; Doutsos et al. 1994; Zelilidis et al. 2002), paleontological dating (e.g., Zygogiannis and Müller, 1982; Barbieri, 1992; Zelilidis et al. 1997), and more recently on hydrocarbon potential and interpretation of seismic data (Kontopoulos et al. 1999; Zelilidis et al. 2002). Some provenance analysis has been made in the sedimentological studies relating to the occurring lithologies and provenance indicators in the sediments, but still no correlation of the sedimentation to the tectonic situation at the surrounding basement rocks has been made. Main aim of this PhD thesis is the examination of the structural data of the Mesohellenic Trough and the bordering basement area in order to understand the structural evolution of the area through time. The recent thermal history of the Pelagonian basement and the provenance of the detrital material in the clastic sediments of the MHT were studied by fission track dating. Apatite and zircon fission track (AFT, ZFT) analysis is applied to samples from the Pelagonian microcontinent along the eastern border of the MHT, and AFT analysis to the sedimentary rocks in the southern MHT, where the whole stratigraphic sequence of the sediment fill is exposed, reaching from the Eocene to the Miocene.

8 Areas near the western border of the trough were also sampled in order to check the provenance of those sediments and/ or any thermal overprint related to tectonic activity. The rapid sedimentation, indicated by the significant thickness of deposits accumulated in the short MHT history, is here examined in means of both the exhumation history of Pelagonian microcontinent (tectonic and erosional denudation) and the subsidence history of the trough. Correlation of the results from the basement and the sediment fill help to unravel the relationship between source and depositional area and to shed light on the evolution of the Pelagonian Zone after the Eocene ( Neo-Hellenic ) orogenic period. FT data from the detrital apatite from the MHT is characterized by two main age populations (i.e. Eocene and Upper Cretaceous to Paleocene ages) and confirm the adjacent to MHT Pelagonian microcontinent as the source of the detrital material. Eocene AFT age populations (between 50 and 30 Ma) in the Eocene (up to Miocene) sedimentary strata indicate a proximal position of the Pelagonian microcontinent, which shows the same or even younger AFT ages. Upper Cretaceous to Paleocene age populations (between ~60 and ~100 Ma) point to a more distant or structurally higher (now eroded) source area. This is concluded from the FT age pattern in the relevant units in the southern F.Y.R.O.M. (Former Yugoslavia, Republic of Macedonia), where a clear AFT age gradient from higher ages in the east to lower ages in the west is documented (Most et al. 2001, Most 2003). The inadequate amount and the bad quality of apatite crystals of the samples in the western part of the trough, led only to an exemplary result. However, the same two main AFT age components (around 40 and 70 Ma) were also recognised in this sample, coming from Upper Oligocene sediments not far from the western MHT margin, showing the Pelagonian microcontinent as the probable, at least partly, source area of those deposits since that time. No correlation with AFT ages of the western bordering basement rocks was available due to inappropriate lithology. The Eocene orogenic event caused only weak thermal overprinting in the rocks of the Pelagonian microcontinent. In its eastern part, the AFT ages show only partial resetting, if any, whereas in its western part the ages were clearly reset during the Eocene event. AFT age-elevation relations, correlation of zircon and apatite FT ages from the same samples, and thermal modelling, based on AFT ages and track length distributions, were all used to reconstruct the low-temperature cooling history of the Pelagonian basement adjacent to the MHT. The results document fast cooling and exhumation in the Eocene that was possibly related to the erosion subsequent of the

9 Eocene thrusting, followed by slow cooling and exhumation during Oligocene and Miocene time. This scenario is confirmed by the AFT data from the detrital material in the MHT sedimentary strata, and the increasing lag times resulted for increasing stratigraphic age. The slow cooling period (between ~30 to 10 Ma) coincides with a stagnation period or crustal extension and possible reheating; this could have been responsible for the partial rejuvenation of the ages of the detrital apatites showed for the oldest (Eocene) formation of the sediment sequence of the MHT. The Mesohellenic Through was characterized as molasse basin by Brunn (1956), but still its strata do not show the expected typical horizontal and undeformed bedding behind an orogenic range. The inclination of the strata, the sequence of the formations in space and time, and the structures characterising the broader region, show that the area experienced a complicated history with different tectonic episodes, which played a significant role to the development of the trough. In this study, a characterization of the MHT is given according to the classification of basin types, after examining the structures and the location where this area can be geodynamically assigned. Numerous structural data, accrued from observations on geometry of kinematics, overprinted criteria, stratigraphic relationships and correlation between various structures, allowed us to distinguish the tectonic events which affected the study area and led to the development of the modern geological situation. These events took place in semi-ductile to brittle conditions from Middle Eocene to Quaternary time. In order to assess the stress regime governing each deformational event, we have calculated its stress tensor from a large amount of fault-slip data. Finally, by the combination of all data and results, a new geodynamic model for the evolution of the MHT area is suggested. The successive tectonic events, in response to which MHT developed, involve isostatic crustal flexure, strike-slip and normal faulting, all related to inferred oblique convergence of the Apulian and Pelagonian microcontinents. This differs from previous interpretations that envisaged foreland flexure related to west-dipping backthrusting (Doutsos et al.1994), or subsidence associated with asymmetrical flexure (Ferrière et al.1998), and/ or normal faulting (Ferrière et al.2004). The MHT evolved geodynamically as a piggyback basin in a foreland setting above westward-emplacing ophiolites and higher Pelagonian units. This agrees as a basin-type characterisation with previous documentations (Wilson 1993, Doutsos et al. 1994, Ferrière et al. 2004). However, great importance is given here to the role of

10 strike-slip faults in the structural evolution of the MHT. Successive stages and changing tectonic regimes recognised in MHT formation are met in strike-slip basins, while experiencing alternating periods of extension and compression. The changing structural settings and repeated episodes of rapid subsidence and uplift, variable depths along the axis of the basin, asymmetry and big length-to-width ratios (4:1), axial infill subparallel to the principal displacement zones, abrupt lateral and vertical facies variations, and of course the presence of strike-slip faults as certainly observed to bound the western side of the MHT are some indicative characteristics of the MHT, typical criteria for the recognition of long-lived strike-slip zones and related basins. As the trough developed due to different tectonic events reported earlier, it corresponds to the pattern of polyhistory strike-slip basins (classification after Busby & Ingersoll 1995). The first stage of basin development, during Middle Upper Eocene, was contemporaneous with the final emplacement of Pindos ocean units and culminated in deformation and uplift of Eocene strata. The Eocene sub-basins developed by crustal flexure and subsidence due to loading of the overthickened Hellenides accretionary prism, associated with a NE-SW transpressional regime and strike-slip faults with reverse component. During the ensuing followed basin closure, intense deformation and uplift in the end of Eocene, the sediments of the first sub-basins were deformed, and placed with a high angle at the western basin margin, locally concordant with the adjacent ophiolitic basement. The Eocene deformation was less intensive at the eastern part. This can be explained by the migration of the main compression towards west through time. The second phase was dominated by strike-slip faults, due to oblique convergence of Apulia and Pelagonian microcontinents. Dextral strike-slip faults of NW-SE to NNW-SSE orientation controlled the subsidence and evolution of the basin from Lower to Upper Oligocene. Strike-slip faults, positive flower structures and rare compressional structures have been developed under a transpressional regime which is characterized only by a decrease in intensity in comparison to the first Eocene regime, and a small shift of the maximum principal stress axis (σ1) from NE- SW towards NNE-SSW. At the end of Oligocene-beginning of Miocene, structural and kinematic evidence disclose a compressional event of local importance, described by a NW-SE trending σ1. Subsidence rates increased in Lower Oligocene in comparison to the Eocene subsidence rates, to decrease again during Upper Oligocene; localized higher rates were evident during the whole period of Oligocene

11 times. These variations are related to the strike-slip tectonic activity and the change of stress regime in the end of Oligocene. The third stage was characterized by low-angle normal faulting along the eastern boundary of MHT during Lower-Middle Miocene, which increased again the subsidence rates at that part of the trough. This associates with the late orogenic collapse and detachment of the Pelagonian nappes. The evolution of the sedimentary basin ended around Middle-Upper Miocene, followed by rapid uplift and marine regression. A compressional event occurred during the late Miocene times, related to reverse and strike-slip faulting. Finally, extensional tectonics affected the area from Upper Miocene to present-day. The Eocene event, as well as the subsequent change in tectonic regime, meets the conclusions deduced from the FT results. Heating of the western Pelagonian microcontinent adjacent to the MHT, during Lower-Middle Eocene, was associated to thrusting, and directly followed by fast cooling and exhumation in Middle-Upper Eocene. The slower cooling and exhumation in the continuing during Oligocene Lower Miocene is associated to the strike-slip faults which cause localized uplift and subsidence (in the area of MHT), while less vertical movements are produced in other places (e.g., the Pelagonian microcontinent). The Miocene extensional period is also intimated from thermal modelling of track length distribution, which indicates a prolonged stay in the same temperature range (or reheating) around ~30 to 10Ma; this can be caused by crustal thinning and rise of the geothermal gradient, accompanying an extensional period. In the latter thermal model, an enhanced uplift is also predicted during the last 10 Ma, which is in consistence with the filling of the basin with sediments and the uplift of the area.

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13 ΠΕΡΙΛΗΨΗ Η Μεσοελληνική Αύλακα (ΜΑ) είναι µία λεκάνη που εκτείνεται στη Βόρεια Ελλάδα και Αλβανία, µε µήκος µεγαλύτερο των 200 χλµ. και πλάτος χιλιοµέτρων. Η λεκάνη είναι συνδεδεµένη µε τις διεργασίες της Αλπικής ορογένεσης. Αναπτύχθηκε από το Μέσο Ηώκαινο έως το Άνω Μειόκαινο, παράλληλα στις Ελληνίδες ζώνες ΒΒ -ΝΝΑ διεύθυνσης, µεταξύ των Εξωτερικών µη µεταµορφωµένων ζωνών (Απούλια µικρο-πλάκα) και των Εσωτερικών µεταµορφωµένων Ελληνίδων ζωνών (Πελαγονικό ηπειρωτικό τέµαχος). Η λεκάνη χαρακτηρίζεται από τη σηµαντική σε έκταση ανάπτυξη, το µεγάλο πάχος των ιζηµατογενών σχηµατισµών (περίπου έως 4 χιλιόµετρα σε κατακόρυφες τοµές), τις πολύπλοκες δοµές, και τις πολυάριθµες µεταβολές στις ιζηµατογενείς φάσεις και τα πάχη των αποθέσεων κατά µήκος και εγκάρσια του άξονα της λεκάνης. Οι ιζηµατογενείς φάσεις περιλαµβάνουν δελταϊκά κροκαλοπαγή, αλλουβιακά κορήµατα, ψαµµίτες και αργίλους υποθαλάσσιων ριπιδίων (τουρβιδίτες), δελταϊκούς και πληµµυρικού πεδίου, και αµµούχα ιζήµατα υφαλοκρηπίδας. Όλα αυτά καθιστούν τη Μεσοελληνική Αύλακα ως τη µεγαλύτερη και πιο σηµαντική λεκάνη του τελευταίου ορογενετικού σταδίου («µολασσικού-τύπου» λεκάνη) των Ελληνίδων. Οι προηγούµενες εργασίες πάνω στη ΜΑ ασχολήθηκαν κυρίως µε τη χαρτογράφηση των σχηµατισµών (π.χ. Brunn, 1956, 1960; Savoyat and Lalechos, 1969, 1972; Mavridis and Matarangas, 1979), την αναγνώριση και ανάλυση των ιζηµατογενών φάσεων (π.χ. Desprairies, 1979; Papanikolaou et al. 1988; Wilson, 1993; Doutsos et al. 1994; Zelilidis et al. 2002), τη χρονολόγηση των σχηµατισµών µέσω παλαιοντολογικών στοιχείων (π.χ. Zygogiannis and Müller 1982; Barbieri 1992; Zelilidis et al. 1997), και σχετικά πιο πρόσφατα µε το δυναµικό υδρογονανθράκων και την ερµηνεία γεωφυσικών δεδοµένων (Kontopoulos et al. 1999; Zelilidis et al. 2002). Έτσι έχει ήδη λάβει χώρα κάποια µελέτη πάνω στην προέλευση των ιζηµάτων της λεκάνης, σχετικά µε τις παρατηρούµενες λιθολογικές φάσεις και ιζηµατολογικούς δείκτες προέλευσης, αλλά παρόλα αυτά, δεν έχει πραγµατοποιηθεί κανένας συσχετισµός µεταξύ της ιζηµατογένεσης και της υφιστάµενης τεκτονικής στα πετρώµατα του υποβάθρου που περιβάλλουν τη λεκάνη. Ο κύριος στόχος της παρούσας διδακτορικής διατριβής είναι η µελέτη των τεκτονικών στοιχείων στην ευρύτερη περιοχή της Μεσοελληνικής Αύλακας, ώστε να κατανοηθεί η τεκτονική της εξέλιξη από το χρόνο σχηµατισµού της λεκάνης έως σήµερα.

14 Η νεότερη θερµική ιστορία του Πελαγονικού τεµάχους που οριοθετεί τη λεκάνη στα ανατολικά, καθώς και η προέλευση των κλαστικών ιζηµάτων της ΜΑ, µελετήθηκαν µε τη µέθοδο θερµοχρονολόγησης των ιχνών σχάσης σε απατίτες και ζιρκόνια (ΙΣΑ και ΙΣΖ). Η ανάλυση των ιχνών σχάσης εφαρµόστηκε σε δείγµατα του Πελαγονικού καλύµµατος κατά µήκος του ανατολικού περιθωρίου της λεκάνης, και στους ιζηµατογενείς σχηµατισµούς στο νότιο τµήµα της λεκάνης όπου αποκαλύπτεται ολόκληρη η στρωµατογραφική ακολουθία των σχηµατισµών της Μεσοελληνικής αύλακας, από το Ηώκαινο έως το Μειόκαινο. ειγµατοληψία πραγµατοποιήθηκε επίσης και στο δυτικό τµήµα της λεκάνης, µε σκοπό τη µελέτη της προέλευσης των ιζηµάτων και σε αυτήν την περιοχή, καθώς και τον έλεγχο κάποιας πιθανής αναθέρµανσης συνδεδεµένης µε τεκτονική δραστηριότητα. Η ταχεία ιζηµατογένεση, τουλάχιστον κατά περιόδους, όπως αυτή υποδεικνύεται από το µεγάλο όγκο των αποθέσεων που συγκεντρώθηκαν κατά την περιορισµένη διάρκεια ανάπτυξης της λεκάνης, εξετάζεται εδώ σε σχέση µε την εκταφή του Πελαγονικού ηπειρωτικού τεµάχους (τεκτονική δραστηριότητα και τη διάβρωση των πετρωµάτων του υποβάθρου) και την ιστορία βύθισης της λεκάνης. Η συσχέτιση µεταξύ των αποτελεσµάτων από τα ΙΣΑ από το υπόβαθρο (Πελαγονική ζώνη) και από το ιζηµατογενές υλικό πλήρωσης της λεκάνης συντέλεσε στον έλεγχο της σχέσης µεταξύ περιοχής προέλευσης και απόθεσης των ιζηµάτων, και έδωσε στοιχεία για την εξέλιξη της Πελαγονικής ζώνης για την περίοδο µετά το Παλαιόκαινο - Ηώκαινο που χαρακτηρίζει την «Ελληνική» ορογένεση. Τα στοιχεία που προκύπτουν από την ανάλυση των ιχνών σχάσης στους απατίτες των ιζηµάτων της ΜΑ (ηλικίας απόθεσης Ηωκαίνου έως και Μειοκαίνου), χαρακτηρίζονται από δύο κύριες οµάδες ηλικιών ΙΣΑ (Ηωκαίνου και Ανώτερου Κρητιδικού-Παλαιοκαίνου), και επιβεβαιώνουν το γειτονικό στη ΜΑ Πελαγονικό ηπειρωτικό τέµαχος ως τη πηγή προέλευσης των κλαστικών υλικών. Η παρουσία ηωκαινικής οµάδας ηλικιών ΙΣΑ (µεταξύ 50 και 30 εκ.χρ.) στα ηωκαινικά (έως και µειοκαινικά) ιζήµατα δείχνει ότι το Πελαγονικό ηπειρωτικό τέµαχος βρισκόταν σε κοντινή απόσταση από τα αποτιθέµενα ιζήµατα, καθότι αυτό παρουσιάζει ίδιες ή και νεότερες ηλικίες απατιτών µε τη συγκεκριµένη µέθοδο. Οι οµάδες παλιότερων ηλικιών ΙΣΑ, του Άνω Κρητιδικού-Παλαιόκαινου (µεταξύ 60 και 100 εκ.χρ.), υποδεικνύουν τροφοδοσία από µία πιο αποµακρυσµένη ή τεκτονικά ανώτερη (τώρα πλέον διαβρωµένη) πηγή. Αυτό συµπεραίνεται από τη διάταξη των σχετικών ηλικιών των αντίστοιχων ενοτήτων στη νότια Π.Γ..Μ. (Πρώην Γιουγκοσλαβική ηµοκρατία Μακεδονίας), που αποτελεί συνέχεια της Πελαγονικής ζώνης και των πιο ανατολικών

15 ζωνών στο Βορρά, όπου παρατηρείται µία ξεκάθαρη διαβάθµιση από µεγαλύτερες ηλικίες ΙΣΑ στα ανατολικά προς µικρότερες ηλικίες στα δυτικά (Most, 2001 et al. Most, 2003). Ο ανεπαρκής αριθµός κρυστάλλων απατιτών καθώς και η κακή τους ποιότητα σε δείγµατα ιζηµάτων και υποβάθρου από το δυτικό τµήµα της Μεσοελληνικής αύλακας, κατέστησαν αδύνατη την ανάλυση των συγκεκριµένων δειγµάτων εκτός ενός, και έτσι οδήγησαν µόνο σε κάποιο υποδειγµατικό αποτέλεσµα. Παρόλα αυτά, οι ίδιες κύριες οµάδες ηλικιών ΙΣΑ (γύρω στα 40 και 70 εκ.χρ.) αναγνωρίστηκαν στα ιζήµατα του Άνω Ολιγοκαίνου κοντά στο δυτικό περιθώριο της λεκάνης, φανερώνοντας έτσι την Πελαγονική ζώνη ως την πιθανή, τουλάχιστον εν µέρει, πηγή προέλευσης και αυτών των ιζηµάτων στο δυτικό τµήµα της Μεσοελληνικής αύλακας από το Άνω Ολιγόκαινο. υστυχώς δεν ήταν δυνατή η εξέταση των ηλικιών ΙΣΑ στο υπόβαθρο του δυτικού περιθωρίου της λεκάνης, και έτσι η συσχέτιση αυτών µε τις γειτονικές ιζηµατογενείς αποθέσεις, λόγω ακατάλληλων πετρολογικών τύπων. Το ορογενετικό γεγονός του Ηωκαίνου προκάλεσε µόνο µία αδύναµη αναθέρµανση στα πετρώµατα του Πελαγονικού ηπειρωτικού τεµάχους. Στο ανατολικό τµήµα αυτού, οι ηλικίες ΙΣΑ φανερώνουν µόνο µία µερική αναθέρµανση, ενώ στο δυτικό του τµήµα οι ηλικίες ΙΣΑ έχουν ξεκάθαρα επαναπροσδιοριστεί κατά το Ηώκαινο. Σχέσεις µεταξύ ηλικιών ΙΣΑ και υψοµέτρου, συσχετίσεις µεταξύ των ηλικιών ΙΣΑ και ΙΣΖ στα ίδια δείγµατα, και θερµικά µοντέλα που βασίζονται στις ηλικίες ΙΣΑ και τις κατανοµές των µηκών των ιχνών σχάσης, χρησιµοποιήθηκαν όλα για την ανάπλαση της θερµικής ιστορίας χαµηλών θερµοκρασιών του Πελαγονικού υποβάθρου που βρίσκεται στο ανατολικό περιθώριο της Μεσοελληνικής Αύλακας. Τα αποτελέσµατα δηλώνουν µία ταχεία ψύξη και εκταφή των πετρωµάτων κατά το Ηώκαινο που πιθανότατα συνδέονταν µε την αυξηµένη διάβρωση που ακολούθησε τις ηωκαινικές επωθήσεις. Μετά ακολούθησε µία αργή ψύξη και εκταφή κατά το Ολιγόκαινο-Μειόκαινο. Αυτό το σενάριο επιβεβαιώνεται από τα δεδοµένα ΙΣΑ στα ιζήµατα της ΜΑ. Η οµάδα ηλικιών ΙΣΑ των 40 εκ.χρ. στο σχηµατισµό του Ηωκαίνου, ίδιες σχεδόν µε τις αντίστοιχες ηλικίες που παρατηρούνται στο Πελαγονικό υπόβαθρο, δείχνουν µία γρήγορη εκταφή και αποκάλυψη των πετρωµάτων στα ανώτερα στρώµατα του φλοιού και συνάµα µία ταχεία διάβρωση και ιζηµατογένεση. Η παρουσία της ίδιας οµάδας ηλικιών ΙΣΑ στους νεότερους σχηµατισµούς αποκαλύπτει την βαθµιαία πιο αργή εκταφή και αποσάθρωση των πετρωµάτων του υποβάθρου. Η περίοδος της αργής ψύξης (µεταξύ 30 και 10 εκ.χρ.) συµπίπτει µε µία περίοδο αδράνειας ή έκτασης του φλοιού, πιθανώς και αναθέρµανσης, η οποία

16 µπορεί επίσης να συνδέεται µε την παρατηρούµενη µερική µετατόπιση των ηλικιών ΙΣ των κρυστάλλων των απατιτών του παλιότερου (ηωκαινικού) σχηµατισµού της ΜΑ προς νεότερες ηλικίες. Η Μεσοελληνική Αύλακα είχε χαρακτηριστεί από τον Brunn (1956) ως «µολασσική» λεκάνη, όµως παρατηρείται ότι τα στρώµατα των αποθέσεών της δεν παρουσιάζονται στην τυπική οριζόντια θέση και χωρίς, ή απλώς µε µία ήπια, παραµόρφωση στο πίσω τµήµα µίας ορογενετικής σειράς, όπως θα ήταν αναµενόµενο. Η κλίση των στρωµάτων, η τοποθέτηση και ακολουθία των ιζηµατογενών σχηµατισµών στο χώρο και χρόνο, και οι τεκτονικές δοµές της ευρύτερης περιοχής, δείχνουν ότι αυτή η λεκάνη χαρακτηρίζεται από µία πολύπλοκη ιστορία µε διαφορετικά τεκτονικά γεγονότα που έπαιξαν σηµαντικό ρόλο στην εξέλιξή της. Στην παρούσα εργασία, δίνεται ένας χαρακτηρισµός της ΜΑ βάσει της ταξινόµησης των λεκανών σε διάφορους τύπους, σύµφωνα µε τις παρατηρούµενες δοµές και της γεωδυναµικής τοποθέτησης του χώρου της λεκάνης. Πολυάριθµα τεκτονικά δεδοµένα που προέκυψαν από παρατηρήσεις πάνω στη γεωµετρία των δοµών, τα κριτήρια αλληλοεπηρεασµού µεταξύ των δοµών, τις στρωµατογραφικές σχέσεις, και την κινηµατική ανάλυση, έκαναν δυνατή τη διάκριση επιµέρους τεκτονικών γεγονότων που επηρέασαν την περιοχή µελέτης και οδήγησαν στην ανάπτυξη της σύγχρονης γεωλογικής δοµής. Τα τεκτονικά γεγονότα έλαβαν χώρα σε ηµι-πλαστικές έως θραυσιγενείς συνθήκες από το Ηώκαινο έως σήµερα. Οι κύριες τάσεις που επικρατούσαν σε κάθε γεγονός παραµόρφωσης υπολογίστηκε από πληθώρα µετρήσεων τεκτονικών γραµµώσεων σε ρηγµάτα. Τελικά, µε τον συνδυασµό όλων των δεδοµένων και των αποτελεσµάτων που συνάχθηκαν, προτείνεται ένα νέο γεωδυναµικό µοντέλο για την εξέλιξη της Μεσοελληνικής Αύλακας. Τα διαδοχικά τεκτονικά γεγονότα που συνδέονται µε την ανάπτυξη της ΜΑ, συµπεριλαµβάνουν ισοστατική κάµψη του φλοιού, οριζόντιες µετατοπίσεις και κανονική ρηγµάτωση, όλα συνδεδεµένα µε τη σύγκλιση, πιθανώς πλάγια, της Απούλιας µικροπλάκας µε το Πελαγονικό ηπειρωτικό τέµαχος. Αυτή η ερµηνεία διαφέρει από τις προγενέστερες που προτάθηκαν, οι οποίες προέβλεψαν µία εµπρόσθια κάµψη συνδεδεµένη µε αντιθετικές επωθήσεις των ανώτερων ενοτήτων προς τα ανατολικά (west-dipping back-thrusting; Doutsos et al.1994), ή βύθιση του

17 χώρου συνδεδεµένη µε ασύµµετρες κάµψεις του φλοιού (Ferrière et al.1998), και/ ή δράση κανονικών ρηγµάτων (Ferrière et al.2004). Η ΜΑ εξελίχθηκε γεωδυναµικά ως µία piggy-back λεκάνη σε µία εµπρόσθια περιοχή, πάνω από τους οφιολίθους και τις ανώτερες Πελαγονικές ενότητες κατά τη διάρκεια της τεκτονικής τοποθέτησης αυτών προς τα δυτικά. Αυτός ο χαρακτηριστικός τύπος λεκάνης συµφωνεί µε προηγούµενες δηµοσιεύσεις (Wilson 1993, Doutsos et al. 1994, Ferrière et al. 2004). Επιπλέον αυτού όµως, στην παρούσα εργασία, ιδιαίτερη σηµασία στην εξέλιξη της Μεσοελληνικής Αύλακας έχει δοθεί στο ρόλο των ρηγµάτων οριζόντιας µετατόπισης. ιαδοχικά στάδια εξέλιξης, που έχουν αναγνωριστεί στη ΜΑ συνδεόµενα µε µεταβαλλόµενα τεκτονικά καθεστώτα, συναντώνται σε λεκάνες οριζόντιας µετατόπισης οι οποίες δοκιµάζουν εναλλασσόµενες περιόδους έκτασης και συµπίεσης. Τυπικά χαρακτηριστικά της Μεσοελληνικής Αύλακας είναι αυτή η µεταβολή των τεκτονικών καθεστώτων µε επαναλαµβανόµενα επεισόδια ταχείας βύθισης και ανύψωσης, τα διαφορετικά βάθη κατά µήκος του άξονα της λεκάνης, η ασυµµετρία της λεκάνης και ο µεγάλος λόγος µήκους προς πλάτους (4:1), η µετατόπιση της ιζηµατογένεσης υποπαράλληλα µε τις κύριες τεκτονικές ζώνες, οι απότοµες πλευρικές και κατακόρυφες µεταβολές των ιζηµατογενών φάσεων, και η οριοθέτηση της λεκάνης από ρήγµατα οριζόντιας µετατόπισης τουλάχιστον στο δυτικό της περιθώριο. Όλα αυτά τα στοιχεία αποτελούν ενδεικτικά κριτήρια για την αναγνώριση διαχρονικών ζωνών οριζόντιας µετατόπισης και των λεκανών συνδεδεµένων µε αυτά. Έτσι, και καθότι η λεκάνη αναπτύχθηκε σε διαδοχικά διαφορετικά τεκτονικά γεγονότα όπως αναφέρθηκε νωρίτερα, πέρα από τον χαρακτηρισµό της ως piggy-back λεκάνη, επίσης καθρεφτίζει το µοντέλο µίας «πολύ-ιστορικής» λεκάνης οριζόντιας µετατόπισης (σύµφωνα µε την ταξινόµηση των Busby & Ingersoll 1995). Το πρώτο στάδιο ανάπτυξης της λεκάνης, κατά το Μέσο-Άνω Ηώκαινο, πραγµατοποιήθηκε ταυτόχρονα µε την τελική τοποθέτηση των ενοτήτων βαθιάς θάλασσας της Πίνδου στις ενότητες των Εξωτερικών Ελληνίδων, και έληξε µε την παραµόρφωση και ανύψωση των ηωκαινικών αποθέσεων. Οι υπο-λεκάνες του Ηωκαίνου αναπτύχθηκαν µε κάµψη και βύθιση του φλοιού εξαιτίας της φόρτισης τoυ πεπαχυσµένου πρίσµατος επαύξησης, και µε τη δράση ρηγµάτων οριζόντιας µετατόπισης µε ανάστροφη συνιστώσα υπό καθεστώς πλάγιας συµπίεσης (transpression). Κατά το επακόλουθο κλείσιµο αυτών των υπολεκανών, της έντονης παραµόρφωσης και ανύψωσης στο τέλος του Ηωκαίνου, τα ιζήµατα των πρώτων υπο-λεκανών παραµορφώθηκαν, και τοποθετήθηκαν µε µεγάλη γωνία κλίσης στο

18 δυτικό τµήµα της Μεσοελληνικής Αύλακας, σύµφωνα µε το παρακείµενο οφιολιθικό υπόβαθρο. Η παραµόρφωση του Ηωκαίνου ήταν λιγότερο έντονη στο ανατολικό τµήµα της αύλακας. Αυτό εξηγείται µε τη µετανάστευση της κύριας συµπίεσης από τα ανατολικά προς τα δυτικά κατά τη διάρκεια δράσης της τριτογενούς ορογενετικής φάσης. Στο δεύτερο στάδιο ανάπτυξης, κυριάρχησε η δράση ρηγµάτων οριζόντιας µετατόπισης, συνδεδεµένη πιθανότατα µε την πλάγια σύγκλιση της Απούλιας και Πελαγονικής µικρο-πλάκας. Η βύθιση και εξέλιξη της λεκάνης κατά τη διάρκεια του Ολιγοκαίνου ελεγχόταν από δεξιόστροφα ρήγµατα οριζόντιας µετατόπισης Β -ΝΑ έως ΒΒ -ΝΝΑ διεύθυνσης. Τα ρήγµατα αυτά, θετικές flower δοµές και συµπιεστικές τεκτονικές δοµές, αναπτύχθηκαν όλα υπό ένα καθεστώς πλάγιας συµπίεσης το οποίο συνεχίστηκε µετά την περίοδο του Ηωκαίνου, αλλά µε µικρότερη ισχύ και συνδεδεµένο µε ένα µέγιστο κύριο άξονα τάσης (σ1) ελαφρώς µετατοπισµένο προς τα ΒΒ -ΝΝΑ σε σχέση µε τη Β -ΝΑ του διεύθυνση κατά το Ηώκαινο. Οι τεκτονικές δοµές µαρτυρούν ένα συµπιεστικό γεγονός τοπικής σηµασίας στα τέλη του Ολιγοκαίνου-αρχές Μειοκαίνου, το οποίο συνδέεται από ένα Β -ΝΑ µέγιστο κύριο άξονα συµπίεσης. Οι ρυθµοί βύθισης δείχνουν µία αύξηση κατά το Κάτω Ολιγόκαινο σε σχέση µε τους ρυθµούς του Ηωκαίνου, και µία επακόλουθη µείωση κατά το Άνω Ολιγόκαινο, ενώ ιδιαίτερα αυξηµένοι ρυθµοί βύθισης έλαβαν χώρα σε συγκεκριµένες θέσεις σε όλη τη διάρκεια του Ολιγοκαίνου. Οι εναλλαγές αυτές στην ταχύτητα βύθισης συνδέονται µε τη δράση των ρηγµάτων οριζόντιας µετατόπισης και την ακόλουθη αλλαγή του καθεστώτος στα τέλη του Ολιγοκαίνου. Το τρίτο στάδιο χαρακτηρίζεται από τη δράση κανονικών ρηγµάτων µικρής γωνίας κλίσης στο ανατολικό περιθώριο της ΜΑ κατά το Κάτω-Μέσο Μειόκαινο, τα οποία αύξησαν και πάλι τους ρυθµούς βύθισης σε αυτό το τµήµα της λεκάνης. Το στάδιο αυτό συνδέεται µε την ορογενετική κατάρρευση και τα ρήγµατα αποκόλλησης στο Πελαγονικό κάλυµµα. Η εξέλιξη της ιζηµατογενούς αύλακας έληξε γύρω στα τέλη του Μέσου-αρχές του Άνω Μειοκαίνου, µε ταχεία ανύψωση της περιοχής και ταυτόχρονη απόσυρση της θάλασσας. Ένα συµπιεστικό γεγονός διαπιστώθηκε κατά τα τέλη του Μειοκαίνου, το οποίο συνδέεται µε τη δράση ρηγµάτων ανάστροφων και οριζόντιας µετατόπισης. Τελικά, εκτατική τεκτονική κυριάρχησε στην περιοχή κατά τα τελευταία εκατοµµύρια χρόνια, από το τέλος του Μειοκαίνου έως σήµερα. Το γεγονός του Ηωκαίνου, όπως και η µετέπειτα αλλαγή στο τεκτονικό καθεστώς, συµφωνεί µε τα συµπεράσµατα που εξήχθησαν από τα αποτελέσµατα των ιχνών σχάσης σε απατίτες και ζιρκόνια. Η θέρµανση του δυτικού Πελαγονικού τεµάχους

19 κατά το Κάτω-Μέσο Ηώκαινο εξαιτίας της συµπιεστικής τεκτονικής και επωθήσεων, ακολουθήθηκε από ταχεία ψύξη και διάβρωση κατά το Μέσο-Άνω Ηώκαινο. Η πιο αργή ψύξη και εκταφή κατά το διάστηµα Ολιγοκαίνου Μειοκαίνου συνδέεται µε τα ρήγµατα οριζόντιας µετατόπισης, τα οποία έχουν ως αποτέλεσµα τοπικές ανυψώσεις και βυθίσεις, ενώ οι κατακόρυφες κινήσεις είναι περιορισµένες σε άλλα σηµεία. Η περίοδος έκτασης του Μειοκαίνου επίσης αποκαλύπτεται από θερµικό µοντέλο εξέλιξης της περιοχής, στο οποίο διακρίνεται µία παρατεταµένη παραµονή στο ίδιο εύρος θερµοκρασιών (ή ακόµα και µία πιθανή αναθέρµανση) µεταξύ των 30 και 10 εκ.χρ., η οποία µπορεί να οφείλεται σε λέπτυνση του φλοιού και άνοδο της ισόθερµων καµπύλων, η οποία συνοδεύει την έκταση. Στο συγκεκριµένο θερµικό µοντέλο προβλέπεται και µία ταχεία ανύψωση κατά τα τελευταία 10 εκ.χρ., η οποία συµφωνεί µε τη γεωλογία της περιοχής, τη σύγχρονη πλήρωση και ανύψωση της ΜΑ κατά το τέλος του Μειοκαίνου.

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21 PREFACE The Mesohellenic Trough (MHT) is a basin that developed during Tertiary times at the area of NW-ern Greece and Albania. Due to its great size and moreover the great thickness of the sedimentary sequences, the Mesohellenic trough is established as maybe the most important late orogenic basin in the Hellenides. The present study is focused on the structural evolution of the Tertiary Mesohellenic Trough. The geometry of deformation in MA and the kinematic analysis of the structures observed in the MA and the basement rocks surrounding the basin led to the distinction of the tectonic events which controlled the development and evolution of the trough. A characterization of the MHT is given according to the classification of basin types and where this area can be geodynamically assigned. At the same time, low temperature thermochronology method of fission track dating has been applied at the University of Tübingen, on apatites and zircons from samples collected from the basement rocks along the eastern border of the MHT and the sedimentary rocks in the southern MHT, in order to unravel the recent thermal history of the basement rocks which supplied the basin with sediments, as well as to examine the relationship between source and depositional area. This thesis was carried out in the frame of the second circle of the PhD postgraduate studies on Structural Geology-Stratigraphy of the Geology Department of the Aristotly University of Thessaloniki and is part of the 03ED375 research project, implemented within the framework of the Reinforcement Programme of Human Research Manpower (PENED), with code 815 and title: Evolution of the stress field and the deformation in the wider area of Greece and generation of strong earthquakes. Contribution in the estimation of seismic risk, and co-financed by National and Community Funds (25% from the Greek Ministry of Development- General Secretariat of Research and Technology and 75% from E.U.-European Social Fund). I m greatly indebted to the State Scholarships Foundation (ΙΚΥ) for financial support during By the completion of this PhD thesis, I would like to express my gratitude to all who helped and supported me for its implementation. First of all I want to address my warmest thanks to my supervisor Professor Mr. Adamantios Kilias, who constituted the corner-stone of the present thesis. Mr. Kilias primarily exhorted me to start my PhD study, and then to go abroad and learn a

22 particular thermochronology method which is not yet applied in Greece, giving me by this way the opportunity to collaborate with another university (University of Tübingen) and to widen my scientific knowledge. I have to notice he was on my side unfailingly from the beginning all through the way to the end, as a teacher and consulter on scientific matters, as a leader pointing out possible directions to follow, and as a friend and an ally showing me trust and comprehension in hours of doubt. Many thanks to the other two Professors of my Three-member Committee Mr. Dimosthenis Mountrakis and Mr. George Migiros for their support and constructive remarks in the final draft, as well as for their significant contribution in previous stage. I also thank deeply the rest members of my Seven-member Board of Enquiry, the Professors Mr. Spyros Pavlides, Mr. Abraam Zelilidis, Mrs. Theodora Rondogianni- Tsiambaou and the lecturer Mr. Markos Tranos, for their time dedicated on the avocation with my thesis and for their meaningful remarks and useful recommendations, which helped me to improve the final version of my thesis. Special thanks to Professor Mr. Zelilidis Abraam for the constructive conversations on sedimentation issues of the Mesohellenic area and his willingness to help any time. I warmly thank the Professor Mrs. Eleftheria Papadimitriou for the help on the financial support as she was the scientifically responsible for the project PENED, for her conscientious attitude, her trust, succour and encouragement until the final fulfillment of the PhD study. Together I would also like to thank my colleagues and friends Dr. Paradeisopoulou Parthena and PhD candidate Domenikos Vamvakaris for their perfect collaboration on the demands of the project and the assumption of many responsibilities. Special thanks to Dr. Annie Rassios from the Institute of Geological and Mineral Exploration (IGME, Kozani department) for her warm interest, her comradeship and support on the progress of science, as well as for the she gave me on every possible matter. I would like to specially thank and to address my gratitude to the Professor Mr. Wolfgang Frisch of the Geology Department of the University of Tübingen for his hospitality at the university, for providing me all the necessary equipment and requisite materials for the preparation of samples and the following analyses, and for his invaluable help during my whole residence in Germany. His contribution has been critical for the progress of my thesis, especially on the part of thermochronology. My warm thanks I also owe to Professor Mrs. Cornelia Spiegel και to the geologist Dr. Martin Danišik for the learning of the fission track technique and their support in any queries. I sincerely thank Mss Dagmar Kost and Dorothea Mühlbayer-Renner for their help on the mineral separation of numerous samples. I also thank warmly my

23 colleagues Dr. John Reinecker and Dr. Christoph Glotzbach for their enlightment on questions raising during the analysis, Dr. Joachim Kuhlemann and Dr. Horst Han for the interesting conversations, and all the rest colleagues in the department for their positive and friendly presence. I couldn t omit to thank the Professor Mr. Wolfgang Siebel for trusting me on research project concerning the analysis of samples and interpretation in the Bavarian Forest of Germany, offering me in this way the chance to work on the geology of a different area and to confirm my knowledge I obtained on the particular method, and on the same time helping me financially to stay longer in Germany after the end of the project on my PhD. I greatly thank my parents Peter and Lina for their financial help for many years, and for staying always beside me, despite their own difficulties, and helping me in any way they could. I also thank my brother Harisis for the patience and comprehension he showed me in hours of great pressure. I want to thank specially me new friends in Germany for the support and love they showed me, contributing to my adjustment there and converting an unknown environment to a familiar one. I also thank deeply my close friends in Greece for their sympathy, comprehension, and their help in any section of my life during this period, as well as my colleagues of the department of Geology in Thessaloniki for their collaboration and support whenever this was needed.

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25 ΠΡΟΛΟΓΟΣ Η Μεσοελληνική Αύλακα (ΜΑ) είναι µία λεκάνη που αναπτύχθηκε κατά τη διάρκεια του Τριτογενούς στην περιοχή της βορειοδυτικής Ελλάδας και Αλβανίας. Λόγω της µεγάλης έκτασης που καταλαµβάνει, καθώς και του µεγάλου πάχους των ιζηµατογενών αποθέσεων, αποτελεί ίσως τη σηµαντικότερη λεκάνη του τελευταίου ορογενετικού σταδίου των Ελληνίδων. Αντικείµενο της παρούσας διδακτορικής διατριβής είναι η µελέτη της τεκτονικής εξέλιξης της Τριτογενούς Μεσοελληνικής Αύλακας. Μελετήθηκαν η γεωµετρία της παραµόρφωσης στη MA και πραγµατοποιήθηκε κινηµατική ανάλυση των τεκτονικών δοµών που παρατηρούνται στη ΜΑ και στους γεωλογικούς σχηµατισµούς του υποβάθρου που περιβάλλουν τη λεκάνη, ώστε να διακριθούν τα τεκτονικά γεγονότα που ελέγχανε τη δηµιουργία και ανάπτυξη της λεκάνης. Επίσης δόθηκε ένας χαρακτηρισµός της ΜΑ βάσει της ταξινόµησης των λεκανών σε διάφορους τύπους, σύµφωνα µε τις παρατηρούµενες δοµές και της γεωδυναµικής τοποθέτησης του χώρου της λεκάνης. Παράλληλα εφαρµόστηκε µέθοδος θερµοχρονολόγησης στο Πανεπιστήµιο του Tübingen της Γερµανίας, σε απατίτες και ζιρκόνια από πετρώµατα του ανατολικού περιθωρίου της λεκάνης και από τους ιζηµατογενείς σχηµατισµούς αυτής, για την εξαγωγή πληροφοριών σχετικά µε την πρόσφατη θερµική ιστορία του υποβάθρου που αποτελούσε πηγή τροφοδοσίας των αποθέσεων της λεκάνης, καθώς και τη σχέση των αποθέσεων µε την ιστορία του υποβάθρου και τις ενότητες προέλευσης των ιζηµάτων. Η διατριβή εκπονήθηκε στο πλαίσιο του δεύτερου κύκλου του Προγράµµατος Μεταπτυχιακών Σπουδών στην ειδίκευση «Τεκτονική-Στρωµατογραφία», του Τµήµατος Γεωλογίας του Αριστοτελείου Πανεπιστηµίου Θεσσαλονίκης και εντάσσεται στο έργο ΠΕΝΕ Ε, µε κωδικό 815 και τίτλο: «Εξέλιξη του πεδίου των τάσεων και της παραµόρφωσης στον ευρύτερο ελληνικό χώρο και γένεση ισχυρών σεισµών. Συµβολή στην εκτίµηση της σεισµικής επικινδυνότητας», για την οικονοµική ενίσχυση για την έρευνα που οδήγησε στη συγγραφή αυτής της διδακτορικής διατριβής, κατά τη διάρκεια των τριών ετών από το 2005 έως το Το έργο συγχρηµατοδοτείται κατά 75% της ηµόσιας απάνης από την Ευρωπαϊκή Ένωση Ευρωπαϊκό Κοινωνικό Ταµείο, κατά 25% της ηµόσιας απάνης από το Ελληνικό ηµόσιο Υπουργείο Ανάπτυξης Γενική Γραµµατεία Έρευνας και Τεχνολογίας, και µε τη χορήγηση γεωλογικών χαρτών του Ινστιτούτου Γεωλογικών και Μεταλλευτικών

26 Ερευνών (ΙΓΜΕ), στο πλαίσιο του Μέτρου 8.3 του Ε.Π. Ανταγωνιστικότητα Γ Κοινοτικό Πλαίσιο Στήριξης. Επίσης οφείλω να κάνω µνεία στο Ίδρυµα Κρατικών Υποτροφιών (ΙΚΥ), που µε ενίσχυσε οικονοµικά για ένα έτος (10/ /2005) πριν την ανάληψη του έργου ΠΕΝΕ. Με την ολοκλήρωση της διδακτορικής αυτής διατριβής θα ήθελα να εκφράσω την ευγνωµοσύνη και τις ευχαριστίες µου σε όλους όσους µου συµπαραστάθηκαν και µε βοήθησαν για την διεκπεραίωση αυτής. Πρωτίστως θέλω να απευθύνω τις θερµές µου ευχαριστίες στον επιβλέποντα καθηγητή µου, κ. Αδαµάντιο Κίλια, ο οποίος αποτέλεσε τον ακρογωνιαίο λίθο της παρούσας διατριβής. Ο κ. Κίλιας µε παρότρυνε να ξεκινήσω την εκπόνηση της παρούσας διδακτορικής διατριβής, ενώ στη συνέχεια µε παρακίνησε και διευθέτησε τον τρόπο για να πάω στο εξωτερικό για εκµάθηση µεθόδου που ακόµα δεν χρησιµοποιείται στην Ελλάδα, δίνοντάς µου έτσι την ευκαιρία να συνεργαστώ µε το ένα ξένο πανεπιστήµιο (Πανεπιστήµιο του Tübingen) και να διευρύνω τις γνώσεις µου στον επιστηµονικό τοµέα. Οφείλω να επισηµάνω ότι υπήρξε δίπλα µου ανελλιπώς σε όλη την πορεία, ως δάσκαλος και σύµβουλος σε επιστηµονικό επίπεδο, ως καθοδηγητής προτείνοντας κατευθύνσεις δράσης, και ως φίλος και σύµµαχος δείχνοντάς µου εµπιστοσύνη και κατανόηση σε στιγµές αµφισβήτησης. Ευχαριστώ θερµά τους καθηγητές κ. ηµοσθένη Μουντράκη και κ. Γεώργιο Μιγκίρο της τριµελούς µου επιτροπής για τη συµπαράσταση και τις εποικοδοµητικές τους παρατηρήσεις στην τελική εργασία, αλλά και τη σηµαντική συµβολή τους σε προηγούµενο στάδιο. Επίσης ευχαριστώ ιδιαιτέρως τα υπόλοιπα µέλη της Επταµελούς Εξεταστικής Επιτροπής, τους καθηγητές κ. Σπύρο Παυλίδη, κ. Αβραάµ Ζεληλίδη, την καθηγήτρια κα Θεοδώρα Ροντογιάννη-Τσιαµπάου και το λέκτορα κ. Μάρκο Τρανό, που αφιέρωσαν το χρόνο τους στην ενασχόληση µε τη διδακτορική µου διατριβή και για τα εποικοδοµητικά τους σχόλια και τις χρήσιµες υποδείξεις τους, βάσει των οποίων βελτιώθηκε σηµαντικά το τελικό αποτέλεσµα της διατριβής. Ιδιαίτερα ευχαριστώ τον καθηγητή κ. Ζεληλίδη Αβραάµ για τις χρήσιµες συζητήσεις πάνω σε θέµατα ιζηµατογένεσης της Μεσοελληνικής αύλακας και την προθυµία του να βοηθήσει ανά πάσα στιγµή. Ευχαριστώ θερµά την καθηγήτρια κα Παπαδηµητρίου Ελευθερία για την οικονοµική υποστήριξη κατά τη διάρκεια εκπόνησης της διδακτορικής διατριβής που ανέλαβε ως επιστηµονικώς υπεύθυνη του προγράµµατος χρηµατοδότησης ΠΕΝΕ, για την εµπιστοσύνη της, την ηθική της συµπαράσταση και την ενθάρρυνση σε όλη

27 τη διάρκεια εκπόνησης της διατριβής. Μαζί ευχαριστώ και τους φίλους και συνεργάτες ρα Παραδεισοπούλου Παρθένα και υποψήφιο διδάκτορα Βαµβακάρη οµένικο για την άψογη συνεργασία τους στις απαιτήσεις του προγράµµατος και την ανάληψη µεγάλο ποσοστό των ευθυνών. Ιδιαίτερα ευχαριστώ τη ρα Annie Rassios του Ινστιτούτου Γεωλογικών και Μεταλλευτικών Ερευνών (ΙΓΜΕ, παράρτηµα Κοζάνης) για τη θερµή της συµπαράσταση, το ενδιαφέρον, τη συναδελφικότητα και την υποστήριξη για την πρόοδο της επιστήµης, καθώς και για τη βοήθεια που µου παρείχε σε ο,τιδήποτε µπορούσε. Θέλω να ευχαριστώ ιδιαιτέρως και να απευθύνω την ευγνωµοσύνη µου στον καθηγητή κ. Wolfgang Frisch του Τµήµατος Γεωλογίας του Πανεπιστηµίου του Tübingen για τη φιλοξενία που µου προσέφερε στο πανεπιστήµιο, µε την ταυτόχρονη προσφορά όλου του απαραίτητου εξοπλισµού και των απαιτούµενων υλικών για την προετοιµασία των δειγµάτων και τις επακόλουθες αναλύσεις, καθώς και για την ανεκτίµητη βοήθειά του καθ όλη τη διάρκεια παραµονής µου στη Γερµανία. Η συµβολή του υπήρξε καίρια για την πρόοδο της διατριβής, ειδικά στο κοµµάτι των θερµοχρονολογήσεων. Θερµές ευχαριστίες οφείλω επίσης στην καθηγήτρια κα Cornelia Spiegel και Γεωλόγο Dr. Martin Danišik για την καθοδήγηση στην εκµάθηση της µεθόδου θερµοχρονολόγησης και για την υποστήριξή τους σε όποια ερωτήµατα προέκυπταν. Ιδιαίτερα ευχαριστώ τις κυρίες Dagmar Kost and Dorothea Mühlbayer- Renner για τη βοήθεια τους στο διαχωρισµό των ορυκτών σε πολυάριθµα δείγµατα. Επιπλέον ευχαριστώ θερµά τους συνεργάτες Dr. John Reinecker και Dr. Christoph Glotzbach για τη διαφώτιση σε ερωτήµατα που προέκυπταν κατά τη διάρκεια των αναλύσεων, τους Dr. Joachim Kuhlemann και Dr. Horst Han για τις ενδιαφέρουσες συζητήσεις, καθώς και όλους τους συνεργάτες του τοµέα για την πρόσχαρη και φιλική τους παρουσία. εν θα µπορούσα να παραλείψω τον καθηγητή κ. Wolfgang Siebel για την εµπιστοσύνη του στην ανάθεση προγράµµατος για την ανάλυση δειγµάτων σε περιοχή της Γερµανίας, προσφέροντάς µου έτσι τη δυνατότητα να ασχοληθώ επιστηµονικά µε τη γεωλογία µίας διαφορετικής περιοχής και να επαληθεύσω τις γνώσεις που απέκτησα πάνω στη συγκεκριµένη µέθοδο θερµοχρονολόγησης, και ταυτόχρονα την οικονοµική ενίσχυση για να παραµείνω επιπλέον διάστηµα στη Γερµανία µετά το τέλος του προγράµµατος χρηµατοδότησης της διδακτορικής έρευνας. Ένα µεγάλο ευχαριστώ οφείλω στους γονείς µου Πέτρο και Λίνα που µε ενίσχυσαν οικονοµικά επί σειρά ετών, και παρά τις δικές τους δυσκολίες ήταν πάντα κοντά µου µε όποιο τρόπο µπορούσαν. Επίσης ευχαριστώ τον αδελφό µου Χαρίση για την υποµονή και κατανόηση που έδειξε σε στιγµές µεγάλης πίεσης.

28 Θέλω επίσης να ευχαριστήσω ιδιαιτέρως τους καινούριους µου φίλους στη Γερµανία για την υποστήριξη και αγάπη που µου έδειξαν, συντελώντας στην προσαρµογή µου εκεί και µετατρέποντας ένα άγνωστο περιβάλλον σε φιλόξενο και οικείο. Επίσης ευχαριστώ βαθύτατα τους κοντινούς µου φίλους στην Ελλάδα για την κατανόηση, τη συµπαράσταση και τη βοήθεια που µου παρείχαν σε όποιο τοµέα µπορούσαν, καθώς και τους συναδέλφους του τοµέα Γεωλογίας για τη συνεργασία και υποστήριξή τους όποτε χρειάστηκε.

29 1. INTRODUCTION In this study, we investigate the development and evolution of the Tertiary Mesohellenic Trough, which is the largest and the most important late orogenic ( molasse-type ) basin in the Hellenides. The MHT is a basin of more than 200 km length and km width, located in northern Greece and Albania (Fig. 1). It developed from Middle Eocene to Upper Miocene time between the external (non-metamorphic) and internal (metamorphic) zones of the Hellenides, on the Tythean ophiolites, following the general trend of the orogen [NNW-SSE] with its long axis. The basin infill reaches a maximum thickness of 4 km in single vertical sections (Zelilidis et al. 2002), while the accumulative thickness of the sediments is much bigger. The different formations show variations in thickness and facies across and along the basin axis; they include fan-delta conglomerates, alluvial fans, turbiditic sandstones and shales, deltaic and flood-plain sandstone and siltstones, and sandy shelf sediments (Zelilidis et al. 1997, 2002). The sequence attracted the interest of scientists and oil companies, because of the favourable conditions for the development of oil or gas reservoirs. Most studies in the MHT focused on mapping (e.g., Brunn 1956, 1960; Savoyat and Lalechos 1969, 1972; Mavridis and Matarangas 1979), sediment and facies analysis (e.g., Desprairies 1979; Papanikolaou et al. 1988; Wilson 1993; Zelilidis et al. 2002), paleontological dating (e.g., Zygogiannis and Müller 1982; Barbieri 1992; Zelilidis et al. 1997), and more recently on hydrocarbon potential and interpretation of seismic data (Kontopoulos et al. 1999; Zelilidis et al. 2002). However, due to the complexity of the basin, lateral facies variations, scarcity of fossils, and partly reworking of the sedimentary material, there still exist discrepancies in published ages and hence lack of precise dating. Furthermore, structural studies were performed towards the assessment of a geodynamic model for the evolution of the basin (Doutsos et al. 1994; Ferrière et al. 1998, 2004; Vamvaka et al. 2004, 2006), but still the proposed scenarios remain controversial. In the end, a paper was carried out very recently in the frame of this PhD thesis, dealing with the low-temperature thermal history of the basement rocks bounding the MHT at the east, and their relation to sedimentary material of the trough (Vamvaka et al. 2009). 1

30 Fig. 1. Geological map showing the main structural domains of the Hellenides in Greece and bordering areas (modified after Kilias et al. 2002). 1. Rhodope (Rh) and Serbomacedonian (Sm) massifs, 2. Pelagonian massif (PI), 3. Attico-Cycladic massif (AC), 4. Axios (/Vardar) zone, 5. External Hellenides (Pa: Parnassos zone, Pi: Pindos zone, G: Gavrovo-Tripoli zone, I: Ionian zone, Px: Paxos zone), 6. Internal HP metamorphic belt, 7. External HP metamorphic belt, 8. Ophiolitic rocks, Sp: SubPelagonian ophiolites, 9. Mesohellenic Trough, 10. Eocene to Miocene thrusts, 11. Paleocene-Eocene thrusts. The geology of the study area is given in more detail in Figure 2. 2

31 The Mesohellenic Through was characterised as molasse basin by Brunn (1956), but still its strata do not show the typical horizontal bedding behind an orogenic range. The inclination of the strata, the sequence of the formations in space and time, and the structures characterising the broader region, show that the area experienced different tectonic episodes, which played a significant role to the development of the trough. Open issues, still existing, triggered our interest to work in the area of Mesohellenic Trough. Some thoughts and questions arising when giving a careful look at the geology of MHT are: The significant thickness of deposits accumulated in the short time of MHT evolution indicates periods of rapid sedimentation. Despite previous sedimentological works, a question remains on how this sedimentation relates to the morphology and the tectonic situation at the surrounding basement rocks area (source area of the sediments). Regarding the different proposed scenarios, a research about the actual main tectonic processes that led to the subsidence of the area and the development of the Mesohellenic Trough seems to be necessary. Searching about the tectonic processes, a question is arised regarding the main stress regime dominating from Eocene till at least Upper Miocene times, and if this was constant through this timeperiod or whether it was subjected important changes. How did the stress regime and its possible changes influence and affect the evolution of the MHT; and therefore which may be a possible geodynamic model for the trough s evolution? Finally, a remaining issue to be clarified, regarding the structures and the location where this area is assigned geodynamically, is as what type of basin should the MHT be considered. The aim of this study is described in the next paragraphs. An overview of the geology of the MHT and the surrounding basement units is given in the continuing. The methods used are described briefly, before the detailed quotation of the observations and results. A thoroughly discussion on the interpretation of data follows, before providing in the end a summation of the final conclusions. 3

32 1.2 Aim of the Study Methods used Main target of this PhD thesis was the study of the structural data of the Mesohellenic Trough and the bordering basement area, in order to understand the structural evolution of the area through time (Fig. 2). The thesis is composed by two parts. The first part of the thesis constitutes a low-thermochronology study on the clastic sediments of the MHT and eastern metamorophic hinterland. We apply fission track (FT) dating techniques to unravel the low-temperature history (associated with the upper crustal levels) of the area that apparently supplied the basin with the huge amount of detrital material. For this, samples were collected from the Pelagonian unit, which borders the MHT to the east. Thermal modelling of the FT data and ageelevation relationships are used to reconstruct the cooling history of the basement rocks. We also sampled the sedimentary formations in the southern part of the MHT in order to define AFT age terrains in the source area of the detrital (clastic) material and to characterize the paleo-source area and its exhumation history in more detail. The western margin of the trough was difficult to be examined due to its inappropriate in most cases lithology and thus the probably inadequate amount and quality of apatite crystals. Still, few sediment samples were indeed collected by the western part of the MHT, in order to check any possible thermal overprint due to tectonic activity. Correlation of the results from the basement and the sediment fill help to unravel the relationship between source area and depositional area and to shed light on the evolution of the Pelagonian Zone after the Eocene ( Neo-Hellenic ) orogenic period. The second part of the thesis is focused on the structural evolution of the basin. Field observations on kinematic indicators, overprinted criteria and sedimentary relationships were made all around the basement bordering areas of the MHT, as well as in the sedimentary formations, aiming to distinguish different tectonic events through time, which led to the development of the modern geological situation. Paleostress analysis was performed, in order to define the alternating stress regime of the trough from Eocene till recent time; finally target is to suggest a new tectonic model for the MHT evolution. 4

33 Summarizing, the targets of the study are: 1. To understand the processes controlling the structural development of the MHT. 2. To provide low temperature thermochronological constraints on the exhumation history of the Pelagonian microcontinent, bordering the MHT to the east. 3. To define the AFT age terrains in the paleo-source area of the detrital material. 4. To investigate if there is any thermal overprint at the sediments of the eastern margin. Also to check out for a possible thermal overprint at the older sediments of the western margin (i.e. adjacent to the contact with the basement rocks), related to tectonic activity. 5. To record the structures of the MHT formations. 6. To define the possible alternating stress regimes at the area of MHT from Eocene till recent time. 7. To construct a geodynamic model for the evolution of the MHT. 5

34 6

35 2. GEOLOGICAL SETTING 2. 1 Geomorphologic features General geological information The Mesohellenic Trough, c. 200 km long by km wide, extends with a NW SE trend from southern Albania in the north through Greece, passing southwards by the cities of Kastoria, Grevena and Kalambaka and finally beneath the younger Neogene and Quaternary deposits of the Thessaly plain (Figs 1, 2). The MHT, as a basin, is characterized by relatively even geomorphologic features, with altitudes varying between 400 and a maximum of 1000 m in the south and 800 to 1750 m in the north. The basement rocks surrounding the trough develop a rather steeper relief, exceeding in places the altitude of 1600 m at the south and 2500 m at the north, and therefore setting out an uplifted rim around the MHT. In the north-western part of the study area (Fig. 2), occurs mountain Grammos, located at the northern-most part of Greece and bordering the north-western part of the MHT and also Greece from Albania. In the continuing to the south, occur the North and South Pindos mountain ranges, with Mountain Koziakas holding the southwestern limit of the basin. Mountains surrounding the MHT at its eastern side, from north to south, are mountain Vernon, Askio, Vourinos and Kamvounia (with summit Vounassa ). Chasia and Voion mountain complexes occur in the south and northwest part of the MHT respectively. Most rivers in the study area follow two main orientations: NW-SE and ENE-WSW. Aliakmonas River is the most significant one; it crosses the trough along strike from the northern part towards the centre, where it curves along the northern margin of Pliocene Karperou Basin and continues with an ENE-WSW to NE-SW orientation to flow finally into the Aegean Sea. Other big rivers in the area of MHT are: Pineios River, which flows with a NW-SE orientation at the south-western part of the trough; Greveniotikos and Venetikos Rivers, which pass through and south of Grevena, respectively, with an E-W to NE-SW orientation, until they join with Aliakmonas R. flux at the centre of the through. 7

36 Fig. 2. Stratigraphic sequences of the MHT and the surrounding basement units (modified after Brunn 1956, and Geological map of Greece of I.G.M.E. 1983). 8

37 Concerning the geological back-round, the Mesohellenic Trough was formed during the latest stages of the Alpine orogenesis and was filled by marine turbidites and siliciclastic shelf deposits. It was developed on the obducted ophiolitic nappe, between the Internal and the External Hellenides. The Hellenides resulted from the collision between the Eurasian plate and the Apulian plate. The collision started in the end of Cretaceous and continued throughout Tertiary times ( Neo-Hellenic orogeny ; Dewey et al.1973, Mountrakis 1986, Jones and Robertson 1991). During this period the Tythean ophiolites were trusted to the west, on to the eastern Apulian margin and constituted the western boundary of the MHT. The Pelagonian microcontinent represents the western Eurasian continental part and borders the MHT to the east. A description of the geological settings of the basement units surrounding the trough is given in the next paragraphs, as well as of the sedimentary formations. The proposed scenarios concerning the development of MHT are also quoted. In the end, a brief review about the basin types classification is provided, in order to distinguish easier which types may be relevant to the case of MHT. 9

38 2.2.1 Pindos Thrust Zone -western boundary of MHT The western side of the Mesohellenic Trough is bordered by the Tythean ophiolites and the trangressively overlying Cretaceous limestones (Fig. 1, 2; Brunn 1956; Jones and Robertson 1991; Mountrakis et al. 1993), which are thrusted together on the Maastricthian to Palaeocene Pindos Flysch and the accompanying oceanic sedimentary assemblage. The pre-orogenic Apulian plate is characterized by continuous carbonate sedimentation throughout Mesozoic times. Continental rifting at this carbonate platform during Upper Triassic resulted in the development of two basins, the so called Ionian and Pindos zones at the western and eastern part of the platform respectively. Shallow marine carbonates deposited on the platform, while deep-sea carbonates and radiolarites deposited in the basins (Karakitsios 1995, Sotiropoulos et al. 2003). Therefore, Pindos zone is considered as a typical basin accommodating a succession of continuous Mesozoic deep-sea sediments, comprising cherts, clayand silt-stones, and deep-water carbonates. The Maastrichtian Paleocene Pindos Flysch deposited at the eastern margin of this basin terminates the sequence. There are two divergent interpretations about the nature of Pindos basin (developed on the eastern part of Apulian platform) and, related to that, the origin of the Tythean ophiolites. The first theory attributes to Pindos basin an oceanic character (e.g. Mountrakis, 1986; Robertson et al. 1996; Dilek and Flower, 2003; Dilek et al. 2005, 2007, 2008; Sharp and Robertson, 2006; Rassios and Dilek, 2008). Subduction of the oceanic crust under Pelagonian microcontinent, started in the Middle-Upper Jurassic, caused the obduction of Pindos Ocean s ophiolite assemblage to the east on the Pelagonian microcontinent during Upper Jurassic. Later, during Paleocene-Eocene (i.e. Neo-Hellenic orogeny time), the Pindos ophiolites were thrusted to the west on the External Hellenides. The second scenario considers the Pindos basin as continental, without the development of an oceanic floor. About the origin of ophiolites, it argues towards the existence of only one oceanic basin in the broader area, Vardar-Axios Ocean, at the east of Pelagonian microcontinent (e.g. Jacobshagen et al. 1978; Hoxha, 2001; Kilias et al. 2001; Gawlick et al. 2007). The ophiolites were thrusted from the east towards west on the Pelagonian block during Upper Jurassic times and further to the west on the External Hellenides during Eocene times. 10

39 Nevertheless the nature of Pindos basin, the whole area was affected by the Tertiary orogenic processes, which were associated with the plate convergence between Apulian and Pelagonian continental blocks, causing compression and crustal thickening. The Tertiary thrusting and nappe-stacking was systematically propagating towards west-southwest (e.g., Jones and Robertson 1991); the latter is also indicated by the constantly younger flysch basins towards the west (Richter 1976; Faupl et al. 1998). The mechanically heterogeneous rock-pile of Pindos zone was highly deformed during the Middle - Upper Eocene, and created a series of imbricate thrust sheets that were emplaced over the Gavrovo-Tripoli zone of External Hellenides (Dercourt 1964; Temple 1968; Fleury 1980, Degnan and Robertson 1998, Xypolias and Doutsos 2000). As a result of the Eocene compression, Pindos zone can be characterized as a large thrust system, extending with a NW-SE strike from Albanides to northern Greece, and continuing to the south, where it is curved at the Peloponnesse in an E-W orientation. Pindos thrust is considered as active until Upper Oligocene times (Fleury 1980, Sotiropoulos et al. 2003). Extension and normal faulting followed the Tertiary compressive tectonics (e.g., Schermer et al. 1993; Sfeikos 1992; Kilias et al. 2002) Pelagonian microcontinent eastern boundary of MHT The Pelagonian microcontinent borders the MHT to the east, and in the study area consists of: (a) a Paleozoic and/or older crystalline basement, which is composed of gneisses and schists and intruded by Upper Paleozoic granitoids; ages of 700 Ma are documented by Anders et al. (2006); (b) a Permian to Triassic volcano-sedimentary sequence and a Triassic to Jurassic platform carbonate cover (Kilias and Mountrakis, 1987; Kilias et al. 1991; Schermer 1993). The Pelagonian microcontinent underwent a complex tectono-metamorphic evolution from Paleozoic to Tertiary times (Yarwood and Dixon, 1977; Jacobshagen, 1986; Kilias and Mountrakis, 1987; Koroneos et al. 1993; Schermer et al. 1993; Most et al. 2001). The wider study area experienced: (1) a Variscan metamorphic event of Carboniferous age under amphibolite facies conditions (Mountrakis 1986; Jacobshagen 1986, Mposkos et al. 2001); (2) an Upper Jurassic to Lower Cretaceous amphibolite- to greenschistfacies tectono-metamorphic event (Eohellenic orogeny) that almost completely 11

40 overprinted the pre-alpine mineral assemblages and textures (Yarwood and Dixon 1977; Barton 1976; Most et al. 2001); (3) an Aptian Albian low-grade retrogressive event (Jacobshagen 1986; Avgerinas et al. 2001; Most et al. 2001; Most 2003); and (4) a Paleocene to Eocene (60-45 Ma) high-pressure/low-temperature metamorphic event found in only basal subunits of the Pelagonian basement, which was overprinted in the very-low grade metamorphic field in Oligocene Miocene times (Schermer et al. 1990, 1993; Kilias et al. 1991; Lips et al. 1998). In contrast, other parts of the Pelagonian microcontinent experienced semi-ductile to brittle deformation in the early Tertiary (Most et al. 2001; Most 2003). The Pelagonian microcontinent was overthrusted by the Tethyan ophiolites and associated sedimentary units in the Middle to Upper Jurassic (Eohellenic orogeny). As described in the previous chapter, according to different interpretations, the ophiolite assemblage either originated from the Vardar-Axios Ocean at the east of the Pelagonian microcontinent, or from two different ocean basins, the Pindos Ocean to the west and the Vardar Ocean to the east. The Early Tertiary (Neo-Hellenic) orogenic event was characterized by WSW-directed nappe-stacking associated with the high-pressure event at the eastern part of Pelagonian block mentioned above. The blueschist unit, exposed in the wider Olympos area (Fig. 1), is sandwiched between the underlying Gavrovo carbonate unit of the External Hellenides (i.e. Olympos unit) and the overlying Pelagonian nappes, and as such marks the suture zone between the Internal and External Hellenides (Godfriaux 1968; Kilias et al. 1991; Schermer et al. 1993; Jolivet et al. 2000). Eocene stacking and crustal thickening (Neo-Hellenic orogeny) were followed by exhumation and crustal-scale extension during the Oligo-Miocene orogenic collapse (Schermer et al. 1990, 1993; Sfeikos et al. 1991; Sfeikos 1992; Kilias et al. 1991, 2002; Dilek et al. 2005). 12

41 The Mesohellenic Trough (MHT) The Mesohellenic Trough developed from the Middle Eocene to the Upper Miocene in the area of the suture located between the Apulian microplate and the Pelagonian continental block. It comprises five, mainly siliciclastic formations (Brunn 1956; Fig. 2), which lie on the Tythean ophiolitic rocks and the overlying Cretaceous limestones. Although more detailed studies have been carried out since the five MHT formations were firstly stratigraphically defined, this division is retained till today (Fig. 2). The later studies involved facies identification, lateral stratigraphical relations and the internal unconformities (e.g. Desprairies 1979; Papanikolaou et al. 1977, 1988; Wilson 1993; Zelilidis et al. 1997, 2002). The biostratigraphy of the basin is based mainly on planktic Foraminifera and nannoplankton (Zygogiannis & Sidiropoulos 1981; Zygogiannis & Müller 1982; Barbieri 1992; Kontopoulos et al. 1999; Zelilidis et al. 2002). The five formations, from bottom to top, are: 1) the Krania Formation (of Middle Upper Eocene age); 2) the Eptachori Formation (of Upper Eocene Lower Oligocene age); 3) the Pentalophos Formation (of Upper Oligocene Lower Miocene age: Chatian- Aquitanian); 4) the Tsotyli Formation (of Lower Middle Miocene age: late Aquitanian Burdigalian (up to Tortonian in places) age); and 5) the Ondria Formation (of Middle Miocene age: Burdigalian - Langian). With the exception of the Krania Formation in the westernmost part of the basin (termed the Gulf of Krania by Brunn 1956), and the south-eastern part of the MHT, the other four formations were deposited parallel to one another through time, from west to east, respectively (Fig. 2). They show an eastward migration within time, as shown by their location and orientation on the map in relation to their age (e.g., Brunn 1956; Zygogiannis & Müller 1982; Barbieri 1992). Accordingly, the youngest formation directly rests on top of the Pelagonian microcontinent along the eastern margin of the trough. At the western edge of the basin, the strata dip towards the ENE at steep angles; dips decrease progressively away from this basin margin, whereas in the centre and along the eastern margin of the basin the strata dip with a low angle towards the WSW. As a result an asymmetrical syncline formed, controlled by structural and 13

42 depositional processes. The MHT splits into two narrower synclines in the south separated by an uplifted structure (Theotokos and Vassiliki villages areas). The main characteristics of the five formations are described below. Surface facies distribution in the MHT and a synthetic stratigraphic column are provided in figures 3A and B, according to Kontopoulos et al. (1999). Krania Formation The Middle to Upper Eocene Krania Formation attains a thickness of at least 1500 m (Brunn 1960; Wilson 1993; Zelilidis et al. 2002); Koumantakis et al. ( Panagia map sheet of I.G.M.E. 1980) estimates a maximum thickness of 2000 to 3000m. It outcrops at the western margin of the MHT ( Gulf of Krania ; Brunn, 1956) and in the southeastern basin near Meteora region. The western occurrence of Krania Formation is characterized by various facies including from base to top coarse breccias, fan-delta roughly bedded mainly ophiolitic conglomerates, shales, and turbitidic fine-grained sandstones and siltstones with nannofossils (Wilson 1993; Zelilidis et al. 2002); they comprise Nummulites, Discocyclina, Actinocyclina, Asterodiscus and Spiroclypeus (Koumantakis et al. 1980). The shaly turbidites compose a sequence at least 1 km thick. Large olistolithic blocks occur in places, especially at the northern margin of the western Krania Formation sub-basin. At the southeastern part, the first deposits are basal conglomerates and Lutetian benthic limestones with Nummulites, Alveolines, Orbitolides and Fabiania, overlain by a succession of marly turbiditic ( flyschmolasse type ) sequences with Globorotalia in interchange with sandstone beds (Upper Eocene; Savoyat et al. 1969, 1972). Eptachori Formation The Uppermost Eocene Lower Oligocene Eptachori Formation occurs along the western side of the MHT and also outcrops in confined areas in the centre of the southern part of the basin (Theotokos and Kalambaka areas, Fig. 2; Brunn 1960; Savoyat et al. 1972; Mavridis et al. 1979). It rests unconformably on the deformed strata of Krania Formation as well as on ophiolites and cretaceous limestones (e.g., Papanikolaou et al. 1988; Doutsos et al. 1994, Ferrière et al. 2004). There is a variation on the estimation of the formation s thickness; in the 1: Geological maps of I.G.M.E. (e.g., Brunn 1960; Savoyat et al. 1972), it is published to be around 1000 m. 14

43 Fig. 3A. Geological map showing surface facies distribution in the Mesohelloenic Trough (Kontopoulos et al. 1999), 15

44 Fig. 3B. Seismic facies interpretation (A1-A2) based on correlation or extrapolating seismic facies to outcrop, and synthetic stratigraphic columns of surface outcrops (B1-B4) and their facies interpretation (C1-C4), based on outcrop data, age determinations and former data by Zelilidis et al. (1997) (Kontopoulos et al. 1999, modified after Zelilidis et al. 1997). B1 and B2 synthetic stratigraphic columns are also provided with colours at the right side of the figure for the easier identification of the MHT formations. 16

45 Tsotyli Formation The Lower to Middle Miocene Tsotyli Formation also reaches a thickness of more than 2000 m and contains Lamellibranchia, Gastropods, Algae and Myogypsina (Mavridis and Matarangas 1979; Mavridis et al. 1985). The base of the formation is characterized by conglomerates that are mainly ophiolite-derived in the northern part of the basin and polygenic, with gneissic conglomerates from the Pelagonian microcontinent, in the south. The conglomerates pass upwards into alternating turbiditic conglomerates, sandstones and shales. According to Zelilidis et al. (1997), these are interpreted as shelf delta deposits, while the finest-grained deposits are related to prodeltaic domains. In the southern part of the basin, the Tsotyli Formation (i.e. upper Meteora conglomerates) lies unconformably on the Pentalophos Formation (i.e. lower Meteora conglomerates), although this unconformity is not observed in the northern part of the basin. In the outer Theotokos village area (Fig. 2), the Tsotyli Formation directly overlies the oligocene Eptachori Formation, whereas in the southernmost part of the MHT, east of Vassiliki village (Fig. 2), it overlies directly the Eocene strata. Ondria Formation The Middle Miocene (Burdigalian Langian) Ondria Formation occurs only in the northern and southern parts of the trough and consists of sandy shelf deposits (i.e. sandstones, marls and limestones; Kumatti et al. 1998). Fossils (e.g., Lithothamnium, Ostrea, Pecten, Clypeaster; Savoyat et al.1971a) in these top MHT deposits divulge that they were accumulated in a shallow-water setting during Middle Miocene times. In the south part, the deposits are estimated to attain a thickness of 600 m (Savoyat et al. 1971a). Ondria formation remains only in a few places of MHT (Fig. 2) probably because of the erosional period that followed (Papanikolaou et al. 1988), but it s also likely that it wasn t deposited all over the eastern part of MHT on top of Tsotyli Formation. The timing of its deposition coincides with a general sea-level drop during the Tortonian (e.g., Haq et al. 1987) and designates the end of the MHT evolution. The shallow-water Ondria Formation may relate to a contemporaneous rather rapid uplift of the basin, without completely filling the basin with clastic material (Papanikolaou et al.1988). 17

46 Except the Eocene Krania Formation, all formations become coarser towards the southern part. This indicates that the southern part where the basin is more narrow, and which is characterized by extensive delta fan deposits (Zelilidis et al. 2002), was likely situated in a more proximal position to the source area than the areas further north. In figure 4 is given a map with iso-depth contours, based on geophysical data (Zelilidis et al.2002). The morphology given for the base of the trough comes to agreement with the sedimentary analysis in the central part of the basin (Zelilidis et al.1997, Avramidis et al.2002), which indicates a shallow deltaic environment of around 50m depth during the end of Oligocene at the area of Kalambaka-Meteora that gets deeper towards north. A shelf environment of c. 200m depth was developed around Theotokos area, while the main basinal region was opening up in the area of Grevena with a depth of c. 700m. Fig. 4. Contours showing the depth where basement rocks occur under the sedimentary formations of the MHT, based on seismic data (depths in meters; v= 4km/sec; modified after Zelilidis et al.2002). 18

47 2.3 Proposed interpretations for MHT evolution Great geodynamic role for the formation and evolution of the Mesohellenic trough is given to the subduction and underthrusting of the External Hellenides towards E- NE (Ferrière et al.2004, Vamvaka et al.2006) and the relative movement of Internal Hellenides towards W-SW (Mountrakis 1986). Besides that, different basin type characterizations, structural and geodynamic interpretations have been considered for the development of the MHT. A study presented from Papanikolaou et al. (1988), associates the beginning of MHT development with tectonic activity related to sliding of big blocks (olistholiths), regional thrusts and reverse faults along the western margin, followed later, around the end of Oligocene, by a gradual tectonic attenuation until the Middle-Upper Miocene. This resulted in the end of sedimentation history of the MHT by the deposition of reef-limestones. This study does not refer extensively to stress regimes or tectonic events, but characterises the trough as a behind-the-arc basin. Basin types that could be characterized as behind-the-arc could probably be the back-arc and retro-arc basins, since the term arc usually refers to the magmatic arc. Still, according to the explanation of the specific authors, with this term they mean a forearc basin, developed between the trench accretionary prism and the volcanic arc, located in North Aegean. Wilson (1993), later, had analyzed the evolution of Krania basin (i.e. the northern Krania basin at the western margin of MHT), which he considered as a separate flysch basin, developed during Middle Eocene on top of the ophiolites. He distinguished Krania basin due to its relation to compression on contrast to the rest formations deposited on top of Krania basin that related exclusively to normal faulting. Wilson supported that Krania basin developed on top of the Tythean ophiolites before their final emplacement over the Pindos flysch, and therefore he characterized it as a piggy-back basin. The first actual structural study on the Mesohellenic Trough was carried out by Doutsos et al. (1994), who suggested a foreland depression developed under compression regime at the eastern side of a huge pop-up structure. The eastern Apulian part was thickened and uplifted, followed later by post-orogenic collapse. The 19

48 trough was formed in front of backthrust faults, dipping to the west, and thus considered as a piggy-back basin. Ferrière et al. (1998) characterized the development of the MHT as a forearc basin at the first stage of its evolution during Eocene, associated with the Pindos basin subduction, while they also consider the basin as a piggy-back basin at the next stages of its evolution, related to the collision of the Gavrovo Tripolitsa platform unit and the following eastward underthrusting of this crustal unit. They suggested it formed in several stages as successively overlapping basins due to asymmetrical flexures that depressed the eastern side of the basin. In a more recent paper, they supported that the subsidence was controlled by normal faulting on the flexures (Ferrière et al. 2004). Zelilidis et al. (2002) studied the sedimentation and stratigraphy of the MHT; they focused on the sedimentary facies, determined their relation to low and high sealevel changes, and compared them to published eustatic curves of sea-level. Analysis of seismic data provided them information for the stratigraphy and the tectonics, and consequently they proposed that the trough developed as a half graben, probably related to strike-slip faults. The present study emphasizes on the significant role of strike-slip faults, while supports the formation of the trough in successive events involving crustal flexuring, strike-slip and normal faulting (Vamvaka et al. 2004, 2006). A multi-phase history of the basin is proposed, started under a transpressional regime during Eocene- Oligocene, followed by extension from the Lower Miocene onwards. 20

49 3. BASIN ANALYSIS BASIN TYPES CLASSIFICATION In this point it was considered important to incorporate a short review about the types of basins that have been distinguished, related to their development and their geodynamic position. This was judged essential, as the present study concerns the structural evolution and geodynamic setting of a large and significant basin. Particular importance has been given in those types of basins, which could interpret the case of MHT, combined with a discussion about previous characterisations and interpretations about the trough s evolution Basins Classification Geologic research on sedimentary rocks takes place in order to find out their geologic history or to evaluate the potential of their economic exploitation. An efficient study requires the combination of all sedimentological, stratigraphical and structural principals. Such a study is usually named as basin analysis, as most of the sedimentary rocks have been deposited in basins. Dickinson (1974) provided the first most intelligible differentiation of basin types related to plate tectonics. Plate tectonics lays major weight on (/examines, emphasizes) the horizontal movements of the lithosphere, which eventually lead to vertical movements due to changes of the crust thickness, the thermal character and isostatic adjustments. The vertical movements account for the subsidence of areas and the development of sedimentary basins, contemporaneously with the uplift of the surrounding areas which feed the basins with detrital material. Main parameters for the basins differentiation and their classification, as well as for their evolution, are: (i) the type of crust (basement rocks) on which the basin set. This can be continental or oceanic or transitional basement from continental to oceanic. (ii) the position of the basin in respect to the margins of the lithospheric plates (i.e. how close is the basin to the plate margins). A basin can develop far away from the tectonic plate margins on a part of the rigid lithosphere, or close to them in mobile zones. (iii) the type of the closest tectonic plate margin, which can be distinguished in the following categories: (I) divergent margins, (II) convergent margins, and (III) transform margins that move laterally. 21

50 According to Wilson cycle of opening and closing of the oceans, which actually constitutes the base of the basins classification, emphasis is given to the plate margins behavior as main criterion of basins differentiation. Cloetingh (1988) has proved the importance of plate margins behavior, by showing that the stresses developed on the plate margins can be transferred in a distance of thousands of kilometers, and therefore control the stress regime and the ways of uplift and subsidence (and thus, the sea level) in the intraplate settings. Plate tectonics settings effect the composition of the strata, and also the structures of the deposits. The carbonate sediments are more usual at mature, divergent margins, due to low relief which dominate in such margins; for this reason there is less detrital material, in contrast to the young rifted margins (either convergent either of lateral movement), which appear high relief (Busby and Ingersoll 1995). The major rivers commonly flow along the axis of tectonic grabens (e.g. Amazon, Mississippi, Rhine) or in internal parallel suture zones (e.g. Ganges, Brahmaputra, Hindus). The petrography of the detrital clastics also depends to a large extend on the regional plate tectonics of the sediments source. Shanmugam and Moiola (1988) supported a wide differentiation between large fan systems at divergent margins where clays dominate, and smaller fans at convergent margins where sands dominate. Finally, tectonics exerts a major control on the rise and fall of sea level, which has a great impact on sedimentary processes and environments and, in turn, on sediment characteristics and stratigraphic relationships (Einsele 2000). After Dickinson s classification that was extensively used, Ingersoll (1988) distinguished 23 types of basins. Later by Busby & Ingersoll (1995) considered that few unusual basins couldn t be classified, while others very common should be differentiated, and therefore proposed 26 types of basins (classification is given in Table 1). As shown in Table 1, these 26 basin types can be sorted in 5 basic categories: Divergent margins basins: they are described by extensional character and they develop on continental crust. They comprise structures of initial continental rifting, which formerly may relate to magmatism or may lead on another occasion to opening of the place where new oceanic floor will develop. 22

51 Divergent Settings Intraplate Settings Convergent Settings Transform (Strike-slip) Settings Terrestrial Rift Valleys Proto-Oceanic Rift Troughs Continental Rises and Terraces Continental Embankments Intracratonic Basins Continental Platforms Active Ocean Basins Oceanic Islands, Aseismic Ridges and Plateaus Dormant Ocean Basins Related to subduction Related to collision Trenches (oceanic) Trench-Slope Basins (oceanic) Forearc Basins (oceanic, transitional) Intra-Arc Basins (oceanic, transitional) Backarc Basins (oceanic, transitional) Retroarc Foreland Basins (continental) Remnant Ocean Basins (oceanic) Peripheral Foreland Basins (continental) Piggyback Basins (continental) Foreland Intermontane Basins (continental) Transtensional Basins (continental and/or oceanic) Transpressional Basins (continental and/or oceanic) Transrotational Basins (continental and/or oceanic) Intracontinental Wrench Basins Hybrid Aulacogens Settings Impactogens Successor Basins Table 1. Basin classification (Busby and Ingersoll, 1995 according to Dickinson s classification, 1974, and Ingersoll 1988). Basins related to long-termed intraplate processes or to the re-activation of ancient weak tectonic margin lines. They are all characterized by extension. Convergent margins basins, which are divided in two categories: (1) Related to subduction and magmatic arcs (fore-, back- and intra- arc basins): some of them develop due to overloading of the overriding crust (oceanic or continental) with sediments, while others (e.g. forearc, backarc) due to extensional stresses (on continental, transitional or oceanic crust). 23

52 (2) Developed during the processes of continental collision: they comprise deep foreland and peripheral depressions, trenches and remnant ocean basins. Basins related to horizontal, lateral plate movements: they may develop on continental or oceanic crust, due to extensional or compressional stress regime. Basins related to hybrid settings, are those which can not be classified in any of the previous occasions. They mainly develop on continental crust, due to far convergent and collisional processes. Several mobile zones have constituted by micro-plates that transferred and collided, sometimes even originating from distal regions; such cases induce difficulties and problems in analysis Subsidence Mechanisms The mechanisms that lead to subsidence and hence basins development, are: (i) crustal thinning due to extension, weathering and/ or magmatic escape. (ii) Mantle lithospheric thickening during cooling that follows either the cessation of extension either the warming due to adiabatic melting or uplift of asthenospheric fluids. (iii) The load from sedimentary or volcanic materials: regional crustal isostatic compensation and regional lithospheric flexure, which depend on the lithosphere s flexibility during sedimentation or volcanic activity. (iv) The tectonic load of crust and the lithosphere: regional crustal isostatic compensation and regional lithospheric flexure, which depend on the flexibility of the underlying lithosphere, during the tectonic emplacement of layers over others (thrusts, reverse faults) or moving of the underlying layers (underpulling). (v) The underground (under crust) load of the crust and the lithosphere: lithospheric flexure during the emplacement of dense lithosphere layers under others (underthrusting). (vi) The asthenospheric flow: dynamic influences of the asthenospheric flow, usually due to the descent of the subducting lithosphere. The crust density increases due to alterating pressure and temperature conditions or/ and the emplacement of higher density fluids in the lower density crust. (vii) The increase of crustal density. 24

53 Fig. 5. Proposed subsidence mechanisms for all types of sedimentary basins (Busby & Ingersoll, 1995). 25

54 Crustal thinning dominates in extensional (divergent) settings. Accordingly, lithospheric thickening is more significant in intraplate settings that originated at divergent plate margins (Fig. 5). Sedimentary load is very important in areas with big sediment supply, especially in those areas where the ocean crust occurs in the continuing of large delta (e.g. continental plateau and remnant ocean basins). Tectonic load prevails in foreland settings (including transpressional). Lithospheric flexure may lead either to subsidence or uplift in far distance from the region where the load is applied. However, all settings bear complicate combinations of several different processes. Additional impacts as paleo-morphology, paleo-climate, paleo-geography and eustatic changes constitute further influences. It should be kept in mind that all basin types represent proto-types of tectonically controlled basins (Einsele, 2000). They offer a starting point for the study and evaluation of basins, but there are no basin types which can be used as a complete model for any other basin (Burchfiel and Royden 1988). 26

55 4. FISSION TRACK THERMOCHRONOLOGY PRINCIPALS AND METHODOLOGY 4.1. Principles of fission track thermochronology Fission tracks are tiny linear zones of damage in the crystal lattice of the host mineral (Fig. 6), created by spontaneous fission when an 238 U atom decays (Wagner, 1968; Fleischer et al., 1975). The so called spontaneous fission tracks are almost exclusively produced by spontaneous fission of 238 U, as other occurring isotopes as 235 U, 232 Th and 244 Pu have much slower rates of decay and therefore produce a negligible number of tracks. Fission tracks are very small (3 to 14 only by enlargement, that may be accomplished by chemical etching, they can be observed under an optical microscope (Fig. 7). The track density depends on the decay rate of spontaneous fission of 238 U, the time of accumulation of fission tracks, and the 238 U concentration in the crystal. At temperatures higher than ~120 C for apatites (e.g., Laslett et al. 1987, Green and Duddy 1989) and ~330 C for zircons (e.g., Tagami and Shimada 1996), fission tracks anneal. When a rock cools to a certain temperature range, known as the Partial Annealing Zone (PAZ), preservation of fission tracks starts. As fission occurs continuously, fission tracks accumulate with time since a rock is located up to a specific depth where temperature is lower than the annealing temperature. Fig. 6. Each fission track is created by spontaneous fission of a single atom of 238 U. 27

56 Fig. 7. Etched spontaneous fission tracks in an apatite crystal Age calculation The method of fission track dating relies on the determination of the relative abundance of the parent (i.e. 238 U) and the daughter product (i.e. spontaneous fission tracks on a given surface of a crystal). The abundance of 238 U can be determined by the number of induced fission tracks created by fission of 235 U in the mica, which is indicative for the 235 U abundance, when the sample is irradiated with thermal neutrons. Given the present-day 235 U/ 238 U ratio of 1/137.88, the 238 U abundance can be calculated. In order to constrain a number of variables such as the exact value of the decay constant of 238 U, the accurate neutron flux and the personal ability of the observer in track recognition, the δ (zeta) calibration factor was introduced (Hurford and Green, 1983). The δ-calibration factor depends on individual differences in the counting procedure. It is determined by analyzing age standards, which can be repeatedly irradiated, together with uranium glass standards [e.g. Corning glasses (CN and U series)]. It can be calculated by the following equation: ( Tstd e 1) s i std G d 28

57 where: δ is the zeta calibration factor (year/cm 2 ); ι α is the decay constant of 238 U; T std is the known age of the standard used; ξ s is the dosimeter (spontaneous) track density from 238 U (track/cm 2 ); ξ i is the neutron-induced track density from 235 U (track/cm 2 ); G is the geometry factor (0.5 for the external detector method, Gleadow and Duddy, 1981; Wagner and Van den haute, 1992). The associated standard error for the δ factor is calculated according to Green (1981) as: ( ) 1 s 1 i 1 d ( std std ) 2 where: N s, N i and N d are the number of spontaneous, induced and glass dosimeter tracks respectively; σ(t std ) and T std are the standard error and the age of the age standard used for calibration. After the δ calibration factor has been determined, the age of each single grain can be calculated by the following equation (Price and Walker, 1963; Naeser, 1967): t s 1 ln s i s d G 1 where: t s is the age of each grain; (ρ s /ρ i ) s is the track density ratio of spontaneous and induced tracks, respectively. The calculation of the sample age is based on the assumption that all grains belong to a single population. Samples from sedimentary rocks may attain more than one age population, but for samples taken from basement rocks (magmatic, metamorphosed), the assumption of a single population is generally justified. However, in order to test the assumption of whether the grains belong to one or more age populations, ρ 2 test is performed on single grain data (Galbraith, 1981). The value of P(ρ 2 ) > 5% (probability of ρ 2 ) is accepted as an evidence of a homogeneous population. Finally, the sample age is presented by an average of the single grain ages. Usually, in order to get a robust result, at least twenty grains are counted in case of basement sample, and at least 50 grains in case of sediment sample. 29

58 4.3. Fission track annealing and modelling Fission tracks shorten with time until they totally disappear, when the rock is remaining in temperatures over a specific range. This process is called annealing and is mainly a function of temperature and time, but other factors as chemical composition of the crystals, crystal structure and pressure play role (Green et al. 1986, Carlson 1990, Wendt et al. 2002, Barbarand et al. 2003). The temperature range, over which the annealing occurs, is called Partial Annealing Zone (PAZ; e.g. Wagner 1979). Over the PAZ, occurs the total annealing zone, where fission tracks are completely annealed in a time of less than 10 6 years. Depending on the temperature profile and the time spent in the PAZ, the fission tracks are progressively shortened. Track-length distributions and FT ages can therefore be used as a basis for modelling the thermal history of the investigated rock (Gleadow et al., 1986a, b; Yamada et al., 1995). The knowledge of zircon PAZ remains rather poor, although several studies were carried out (e.g., Hasebe et al. 1994, Yamada et al. 1995a, b). Results suggest a temperature range for zircon PAZ between ~ C, with a mean effective closure temperature of ~240 ± 50 C for cooling rates of 10 C/Myr (e.g., Tagami and Shimada, 1996). For apatites, the PAZ is generally defined as a temperature range between 60 and 120 C, with mean effective closure temperature of 110 ± 10 C (e.g., Laslett et al. 1987, Green and Duddy 1989). Chemical composition of apatite (generally written as Ca 5 (PO 4 ) 3 [F,Cl,OH]) influences the thermal retentivity of the fission tracks (e.g., Green et al. 1989). The relative proportion of Cl and F are accepted to play important role in the annealing process (i.e. Cl-apatite is more resistant to annealing than F- apatite; e.g., Gleadow and Duddy 1981, Green et al. 1986, Donelick 1991), but still little is known about the other ions for apatites with different composition, despite the numerous studies. Another factor that controls the annealing of apatites is the track orientation with respect to the crystallographic c-axis. Tracks at higher angles to the c-axis anneal faster than tracks at lower angles (Green et al. 1986, Donelick et al., 1990, Crowley et al. 1991). Dpar value (i.e. the etch pit diameter of fission track parallel to the crystallographic c-axis at its intersection with the polished, etched and analyzed apatite surface; Crowley et al. 1991, Naeser 1992, Burtner et al. 1994) is regarded as the best 30

59 parameter to characterise the effect of the apatites chemical composition on annealing (Ketcham et al. 1999). It s found out that the solubility of apatite is dependent upon the chemical composition in such a way that the fluorine-rich apatites are more soluble than the chlorine-rich apatites (Donelick et al. 1999). Instead of the chemical composition, the solubility, estimated by measuring the maximum etch pit diameter, can serve as an equal annealing kinetic parameter. A small Dpar value ( κm) is representative feature of F-apatites and faster annealing, while larger Dpar value ( κm) stands for Cl-apatites and slower annealing. Fission track modelling As annealing of fission tracks occurs by the increase of temperature, track lengths shorten and thus track density reduces. In contrast, when rocks are passing at levels higher than the APAZ (i.e. the upper crust levels), then annealing of fission tracks stops and thus track density increases. Therefore, track density can be used for the determination of the apparent age of a sample, which is indicative for the last time the apatite experienced temperatures within the APAZ. On the other hand, track length distributions, with respect to the age of the sample, signify how fast cooling occurred, and consequently they can be used for obtaining information about the thermal history of the sample. Confined tracks are fully enclosed in the crystal and etched via cleavage planes or cracks (called as TINCLE s) and via surface tracks (called as TINT s; Fig. 8). For undisturbed basement, a mean confined track length distribution of about 12 to 14 κm is representative for slow cooling or regional metamorphism, while rapidly cooled rocks (usually intrusive and volcanic rocks) display a mean track length distribution of 14 to 15.5 κm; bimodal and mixed distributions are attributed to different types of two-stage thermal overprinting (Gleadow et al. 1986). In general, the lack of shortened tracks indicates a short stay in the PAZ and therefore a fast cooling of the rock. In contrast, the existence of shortened tracks reveal a long stay in the PAZ and thus a slow cooling or even a stagnation period of the rock within the temperature range of PAZ. 31

60 Fig. 8. Etched spontaneous fission tracks in an apatite crystal. Etch pits are parallel to c-axis. Confined tracks and the dpar (etch pit) are pointed out on the picture. Several programs exist, which enable to model the time-temperature (tt) path of a sample. For this study, the HeFTy program by R. Ketcham (2005) was used with the multi-kinetic annealing model of Ketcham et al. (2007 b), with c-axis projected confined track length data (2007a) and Dpar values as kinetic parameter for the thermal history modelling. The program defines a number of tt paths, passing statistical criteria and confirming to user-entered constraints, that best reproduce the measured data. It provides a graphic display of the modelling results, showing the best fit evolution trend and statistically acceptable limits. The best fit thermal history may neither be geologically meaningful nor unique. In order to constrain the significance of the model results other geological data like metamorphic events or structural data are necessary to decide which thermal history fits best the geological situation of the region. 32

61 5. FISSION TRACK ANALYSIS IN MESOHELLENIC TROUGH AND THE ADJACENT PELAGONIAN BASEMENT 5.1. Sampling and Analytical Procedure We focused our sampling campaign on the Pelagonian basement to the east of the MHT, which contains appropriate lithologies to apply the method (Fig. 9). No FT ages have been reported from this area so far. Sampling was performed in two areas close to the southern and the northern parts of the MHT. The sampling area is cut by normal faults but otherwise forms a tectonically homogeneous block. The landscape is hilly, and therefore no steep profiles to study age-elevation relations were available. Samples were also taken from sandstones (Krania and Eptachori Fm.) and conglomerates (Pentalophos and Tsotyli Fm.) from the southern part of the MHT in order to investigate their provenance, the relation between depocenter and exhumation area, and a possible thermal overprint of the sediments. In the sampling area the whole stratigraphic sequence of the sediment fill is exposed, reaching from the Eocene to the Miocene (Fig. 9). The proximity and direct contact to the Pelagonian basement as well as the pebble compositions clearly mark the basement to the immediate east of the MHT as the most probable source area of the detrital material. Few samples (five sediment samples and one from the ophiolitic basement) were also collected from the western part of the basin (Fig. 9), although the petrological types of the ophiolitic basement rocks were rather inappropriate for FT analysis as they are not indicative for containing adequate amount of apatite crystals. The sediment samples were taken mainly from Krania and Eptachori Formations, near their contact with the basement rocks and also a little further, in order to check for a thermal overprint related to tectonic activity. A possible overprint in sediments close to tectonic zones, could provide information about the specific time that the observed tectonic structures were developed. This would be an additional confirmation on the conclusions deduced by the structural observations, which will be discussed in the second part of this study. 33

62 Fig. 9. The samples taken from the whole area of the MHT and the surrounding basement rocks (yellow points). (Map modified after Brunn 1956, and Geological map of Greece of I.G.M.E ) Apatite and zircon crystals were separated using standard magnetic and heavy liquid separation techniques. Apatites were embedded in epoxy and zircons in PFA Teflon ΣΜ. Prepared mounts with grains were grinded and polished. Spontaneous tracks in apatites were revealed by etching with 5 M HNO 3 solution for 20 seconds at 20 ºC. Zircons were etched in an eutectic mixture of KOH and NaOH at 215 ºC for 20 34

63 to 80 hours (Zaun and Wagner 1985). For FT analysis the external detector method (Gleadow, 1981) was used. FT ages were calculated using the δ (zeta) age calibration method (Hurford and Green 1983). Samples were irradiated in the thermal column of the FRM-II nuclear reactor in Garching (TU München, Germany). Neutron fluence was monitored using the Corning glass dosimeters CN-2 (for zircons) and CN-5 (for apatites), with known uranium content of 37 ppm and 12 ppm, respectively (Hurford and Green 1982). After irradiation, Goodfellow mica ΣΜ detectors were etched in 40% HF for 30 minutes at 21 ºC. Finally, the mounts with corresponding micas were attached side by side on a glass slide. Only crystals with a well-polished surface parallel to the crystallographic c axis and homogenous uranium distribution were analysed. A minimum of 25 and 60 grains from each sample from basement and sediments, respectively, was counted. For the sediment samples, which are characterised by low ρ 2 (chi-square) values, the PopShare program (Dunkl and Székely 2003) was used, in order to deconvolute the data into different age groups, using the best fitting solutions. As an additional control of our evaluation, we also used BinomFit Program (Brandon 1992). Criteria for the statistical soundness of our deconvolution were the minimized values of goodness of fitting RMS (root mean square) for the PopShare program (Dunkl and Székely 2003), and the F-test for the BinomFit program (Brandon 1992). AFT data were used to predict time-temperature (tt) paths, which were subsequently taken for thermal history inversion of individual samples. Modelling was carried out with the HeFTy program of Ketcham (2005), based on the multikinetic annealing model of Ketcham et al. (2007b) and with c axis projected track length data (Ketcham et al. 2007a). tt paths were statistically evaluated and categorised by a value of goodness of fit (GOF), in which a good result corresponds to a value of 0.5, an acceptable result corresponds to a value of 0.05, and a GOF of 1 is the optimum. For details, I refer the reader to Ketcham (2005). The input parameters for each sample used in this study are its central FT age with 1ζ error, the c axis projected track length distribution and, as a kinetic parameter, the Dpar value (e.g., Burtner et al. 1994). (A more detailed description about the analytical procedure of the preparation of the samples, the age determination and thermal modelling, is provided in Appendix A). 35

64 5.2. Thermochronological Data ZFT and AFT ages were published from the Pelagonian microcontinent and the adjacent parts of the Axios (/Vardar) Zone in northernmost Greece (region of Macedonia) and the southern (former Yugoslav) Republic of Macedonia (F.Y.R.O.M.; Fig. 10; Most et al. 2001; Most 2003). ZFT ages range from 86 to 39 Ma. Eocene ZFT ages (53-39 Ma) are found west of the Florina-Prilep Neogene depression. In the Pelagonian and Axios Zone to the east of this depression the ZFT ages are Upper Cretaceous to Paleocene (86-58 Ma), but show no regional trend. Fig 10. Apatite and zircon FT ages in F.Y.R.O.M. (former Yugoslav Republic of Macedonia) and northern Greece (ages by Most et al. 2001; Most 2003). Western part of this area is the northern continuation of the present study area (modified after Vamvaka et al. 2009). AFT ages range from 60 to 25 Ma (Fig. 10; Most et al. 2001, Most 2003). They are grouped in orogen-parallel zones and show a general and gradual trend from higher ages in the east to lower ages in the west (i.e., the ages get younger towards 36

65 more external positions in the orogen). This age trend is in agreement with AFT ages between 21 to 11 Ma reported from eastern Albania (Mućčeku et al. 2006), since they derive from a still more westerly (external) position. Most (2003) also performed apatite fission track length measurements, which allowed thermal modelling of his samples, preferentially from the eastern part of his study area according to the availability of data. The modelled thermal history displays slow continuous cooling between 80 and 50 Ma, followed by remarkably faster cooling up to 40 Ma or a little later. Additionally, Most (2003) provided few AFT data from the southern part of the Mount Olympos region to the east of our study area (Fig. 1). Apatite ages of Ma indicate Low to Middle Miocene exhumation related to doming in the tectonic windows of the Olympos and Ossa Mountains. 37

66 5.3. Results More than 50 samples have been prepared in mounts and irradiated to be ready for AFT analysis (Fig. 9). The large number of the samples prepared was aiming to avoid problems arising by the quality of the apatite crystals and the further possibility to apply the method without having a great risk to receive unreliable results. Excluding those samples which were impossible to count, 26 were chosen between the rest for analysis, according to their relative spatial location and their altitude (in order to examine in the best possible way the different structural levels), and for the sediment samples the formation they belonged to (Fig. 11). Unfortunately, the samples collected from the western part of the MHT, were all but one impossible to count due to the lack of adequate number of apatite crystals and additionally of the very bad quality of the few apatites. The countable sample came from the northern part, from Pentalophos Formation (i.e., Ept-2; see figures 9 and 11), while the second sample from the same region, located closer to the MHT basement boundary and of younger age (Lower Oligocene), was also uncountable. AFT analyses were performed on 17 samples from the Pelagonian basement (5 from the northern and 12 from the southern part), and 9 samples from the MHT sediment fill, from which one belongs to the western part and the other 8 were taken from the southern part of the trough where the whole stratigraphic sequence is exposed (Figs 12, 13; Tables 2 and 3). ZFT dating was applied on 4 samples from the basement (two samples from northern and two from southern Pelagonian basement rocks; Fig. 12 and Table 2) Pelagonian microcontinent samples (western part adjacent to MHT) All FT ages from the basement are displayed as central ages with errors of ±1ζ (Fig. 12, Table 2), and all passed the ρ 2 test at the 5% level. AFT ages in the southern basement study area range from 44 to 25 Ma, while the ZFT ages are 45 and 47 Ma. Samples from the northern basement study area yield AFT ages between 43 and 32 Ma and ZFT ages of 48 and 50 Ma. In general, as can be observed in figure 12, both areas show very similar age patterns. 38

67 Fig. 11. The north (A) and the south (B) sampling areas. Also the sediment sample Ept-2 at the western part is indicated in a separate frame (C). (Map modified after Brunn 1956, and Geological map of Greece of I.G.M.E ) 39

68 The Dpar values are rather constant and thus indicate that the annealing kinetic parameters of the different apatites are similar (Table 2). In the majority of cases, single Dpar values show a narrow scatter in the overall range between 1.3 and 2.7 µm, which indicates low resistance against annealing. Track-length distributions were measured in four basement samples with different AFT ages (Table 2 and Fig. 12). The distributions of samples AV-K2, AV-18B and AV-44 are narrow with mean confined track lengths and standard deviations (MTL ± SD) of ± 1.3, ± 1.60 and ± 1.82 µm, respectively. Sample AV-23 displays a wider track-length distribution, which ranges from 7.5 to 16.5 µm and shows a mean value of ± 2.08 µm. The longer MTL (>13.5 µm) and shorter SD, typical for rapid cooling through the APAZ, are found in the two older samples (AV-K2 and AV-18B, ~42 Ma). In contrast, the younger samples (AV-44, AV-23) are characterised by shorter MTL and longer SD, indicating slower cooling or even prolonged stay in the APAZ Sediment samples of the Mesohellenic Trough All sediment samples display a big range of AFT single-grain ages, from 21 to 120 Ma, common for sediments since material is usually derived from more than one lithologic units. Half of the sediment samples at the southern part of Mesohellenic Trough, and also sample Ept-2 from Pentalophos Formation from the western side of the trough, display low ρ 2 probability values and thus indicate more than one age populations (Table 3). Four of the samples passed the ρ 2 probability test, displaying central ages between 44 and 56 Ma. The mean Dpar values and the range of single Dpar values of the most samples resemble to those of the basement samples (ranging in overall between 1.3 and 2.7 µm), showing in general low resistance against annealing (Table 3). In two sediment samples (AV-81 and AV-52) the Dpar values show an average of 3.2 µm. However, no distinct correlation between Dpar values and ages is observed. 40

69 Fig. 12. Locations of samples, FT ages, and track-length distributions for the two parts of the study area (A and B); Sample Ept-2 from the western part is shown at the corner of the north area. For location of maps, see Figure 11. Central FT ages ± 1ζ error are shown for the Pelagonian basement samples; in the sediment samples from MHT, the youngest AFT age component ± 1ζ is given. 41

70 Here it was considered meaningful to give some more detailed information for sample AV-81. The apatite crystals were separated, prepared and analysed in two different mounts, as they were characterised by two groups of different size, of clearly bigger and smaller crystals. The mount with the small crystals (i.e. AV-81-s, see Table 3) was of a very bad quality and only 19 crystals were counted, while 61 crystals were analysed for the mount with the bigger crystals (i.e. AV-81-b; Table 3). The sample AV-81-s passed the ρ 2 probability test and showed an AFT agepopulation of ~40Ma and an average of Dpar values of 2.7 κm. The results from sample AV-81-b indicated also a single age-population with a central AFT age of ~47 Ma, while the Dpar values from this mount showed a higher average of 3.4 κm. Finally, the results from both mounts of sample AV-81 were integrated and they are discussed further in interpretation. 42

71 Fig. 13. Stratigraphic sequence in the southern MHT. Positions of sample sites are also shown. 43

72 Table 2. AFT and ZFT data for samples from the Pelagonian microcontinent a. Sample apatite data Geological unit Pelagonian Unit (P) y x Alt. (m) n ρ s N s ρ i N i ρ d N d P(χ 2 ) [%] D Central age ±1ζ [Ma] MTL±s.d. [κm] Dpar±s.d. [κm] U [ppm] AV-K2 AV-K4 AV-K6 AV-K9 AV-K10 AV-7 AV-27 AV-18A AV-18B AV-19 AV-29 AV-33 AV-37 AV-38 AV-40 AV-44 AV-23 P, North P, North P, North P, North P, North P, South P, South P, South P, South P, South P, South P, South P, South P, South P, South P, South P, South ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± Zircon data AV-K4 AV-K10 AV-40 AV-27 P, North P, North P, South P, South ± ± ± ± a x, y: Geographic coordinates (Lat/ Long), WGS84 Datum; Alt: Altidute; n: number of counted grains; ξ s / ρ s : spontaneous/ induced track densities respectively (10 5 tracks.cm -2 ); N s / N i : number of counted spontaneous/ induced tracks; ρ d : dosimenter track density (10 5 tracks.cm -2 ); N d : number of tracks counted in docimeter; P(ρ 2 ): probability obtaining Chi-square value (ρ 2 ) for n degree of freedom (where n is number of crystals minus 1); D: Dispersion; the central age is given, where the ±1ζ stands for the ±1 standard error (Galbraith and Laslett, 1993); MTL: mean horizontal confined track length; Dpar: the mean etch pit diameter of fission tracks, where each etch pit diameter was averaged from 4 measurements per analysed grain. The ages were calculated using zeta calibration method (Hurford and Green, 1983), glass dosimeter CN-5 and zeta value of ±8.25 yr.cm -2 for apatites, and glass dosimeter CN-2 and zeta value of 123.6±2.1 yr.cm -2 for zircons. 44

73 Table 3. AFT data for samples from the Mesohellenic Trough b. Sample Formation (Fm.) y x Alt. (m) Apatite data MHT AV-80 Krania Fm n ρ s N s ρ i N i ρ d N d P(χ 2 ) [%] D Singlegrain AFT ages [Ma] 1st AFT agepopulation ±1s.d. [Ma] ± ±2.7 2nd AFT agepopulation ±s.d. [Ma] AV-81-s Krania Fm ± ± ±6.2 AV-81-b Krania Fm ± ± ±15.1 Dpar ±s.d. [κm] U [ppm] 66.8± ± ± ± Av-81-all ± ± ± ± ±3.5 AV-93 Eptachori ± ± ± Fm. AV-94 Eptachori ± ± ± Fm. EPT-2 Pentalophos ± ± ± Fm. / West AV-50A Pentalophos ± ± Fm. 47.4± ±11.5 AV-52 Pentalophos ± ± Fm. 37.9± ±9.4 AV-101 Tsotyli Fm ± ± ± AV-107 Tsotyli Fm ± ± ± ±5.8 b x, y Geographic coordinates (Lat/ Long), WGS84 Datum; Alt: Altidute; n: number of counted grains; ξ s / ρ s : spontaneous/ induced track densities respectively (10 5 tracks.cm -2 ); N s / N i : number of counted spontaneous/ induced tracks; ρ d : dosimenter track density (10 5 tracks.cm -2 ); N d : number of tracks counted in docimeter; P(ρ 2 ): probability obtaining Chi-square value (ρ 2 ) for n degree of freedom (where n is number of crystals minus 1); D: Dispersion; the central ages are given for the sediment samples with high P(ρ 2 ), where the ±1ζ stands for the ±1 standard error (Galbraith and Laslett, 1993), the range of single-grain AFT ages, and the first and second AFT age-populations for all the sediments samples ± standard deviation, as calculated with PopShare (Dunkl and Székely, 2003); Dpar: the mean etch pit diameter of fission tracks, where each etch pit diameter was averaged from 4 measurements per analysed grain. The ages were calculated using zeta calibration method (Hurford and Green, 1983), glass dosimeter CN-5 and zeta value of ±8.25 yr.cm

74 5.4. Interpretation and Discussion Based on our AFT and ZFT data, and taking previous results from a wider area into consideration, we arrive at conclusions regarding the exhumation history of the Pelagonian microcontinent during the evolution of the MHT, the AFT age terrains of the paleo-source area of the detrital (clastic) material, and the possible resetting of the AFT ages of the sediments. The interpretations given are discussed in this chapter. However, firstly an evaluation of the FT data of Most et al. (2001) and Most (2003) from the Pelagonian microcontinent in the southern F.Y.R.O.M. and adjacent northern Greece (Fig. 10) was considered necessary, because their conclusions are of importance for evaluating the exhumation history of the source area of the clastic sediments in the MHT in the study area Pelagonian microcontinent in the southern F.Y.R.O.M. In the southern Republic of Macedonia and northern Greece the pre-eocene ZFT ages of the eastern Pelagonian and the Axios Zones (Fig. 10) reflect cooling after Upper Jurassic to Lower Cretaceous (Eohellenic) metamorphism in the temperature range below 300 C, and/or heating and partial rejuvenation during the Eocene orogenic event. The Eocene ZFT ages in the western part of the Pelagonian microcontinent (west of the Florina-Prilep depression) partly coincide with the climax of the Eocene orogenic event around 50 Ma. The slightly younger ZFT ages mark cooling shortly after the thermal peak (Fig. 10). AFT data show that in the eastern parts of the Axios Zone, the rocks cooled below 100 C already by Paleocene time, whereas there is no evidence of the Eocene event affecting the AFT system in this area (Figs 10 and 14). Moving westward, the AFT ages get gradually younger. In the western part of the Axios Zone and the easternmost part of the Pelagonian microcontinent, the AFT ages around 52 to 45 Ma coincide with the peak of Lower to Middle Eocene heating, which therefore must have reached temperatures of slightly more than 100 C. 46

75 Fig. 14. Temperatures of the basement rocks of the present erosion level at around 40 to 45 Ma, resulting from the FT ages of the area (see also figure 10). The fission track age patterns from the southern part of F.Y.R.O.M. show that in the present erosion level, temperatures during the Eocene orogenic event were higher in the western part than in the eastern part (Fig. 14). West of the Florina-Prilep depression, temperatures of >250 C were reached, as evidenced by the ZFT ages. Accordingly, in the Pelagonian microcontinent further east and the adjacent western part of the Axios Zone, temperatures were lower during the Eocene, as shown by the ZFT ages that are older than Eocene. A gradual decrease in temperature toward the east is denoted by the increasing AFT ages. The age pattern shows a clear trend to higher denudation values toward the west after the Eocene orogenic event. A local heat source in the western part could be excluded by the gradual westward decrease of the ages with a low gradient (and continuing into Albania; see above). An alternative interpretation may be that thrusting in the Axios Zone had already started in the Paleocene and had propagated westward until the Eocene. Thermal history modelling displays slow cooling between 80 and 50 Ma and faster cooling between 50 and 40 Ma or a little later (Most 2003). We consider this increase in cooling (and hence in exhumation) rate as a time constraint for Eohellenic orogenic movements and subsequent denudation in Lower to Middle Eocene time. The Eocene orogeny probably mainly resulted in reactivation of older compressional structures in this region. A significant decrease of the cooling rate or even prolong stay of some samples in the APAZ characterises the post- Eocene period. 47

76 Pelagonian microcontinent in the study area (NW Greece) The Pelagonian microcontinent in the study area at the eastern border of the MHT shows the same ZFT ages as in the area immediately to the north, west of the Florina-Prilep depression. This means that the Pelagonian microcontinent in the study area was heated to temperatures high enough (i.e., ~250 C or higher) to reset the ZFT ages in the Eocene. The AFT ages are also the same as those further to the north, i.e. in the areas both west and east of the Florina-Prilep depression. As a result, both northern and southern regions can be considered as a continuous zone with the same thermal history, which is also backed by their spatial relationship (Fig. 1). In the Pelagonian microcontinent along the eastern border of the MHT, the AFT data display a general positive age-elevation relationship (Figs 15, 16A; Table 2). Two profiles show similar AFT ages and age-elevation patterns (Fig. 15), although horizontal distances of the samples in the profiles are rather large (10 and 8 km). In both profiles the peak sample shows a clearly higher age (43 and 41 Ma), whereas the samples from lower elevations are around 35 to 30 Ma and overlap within limits of error. In both profiles, the relatively young AFT ages were measured up to high elevations, so that a jump in age occurs in the proximity of the peak sample. The only feasible explanation for this pattern is that in both cases normal faults separate the peak areas from the slopes. Such faults correspond to the overall tectonic pattern of the area, which experienced extensional tectonics after Eocene-Oligocene time (Schermer et al. 1993; Kilias et al. 2002; Vamvaka et al. 2006). The faults preclude the calculation of exhumation rates in the two profiles. However, the age-elevation relationships of all AFT ages in the whole study area, show strong clustering of ages around 40 Ma (Fig. 16A). This observation suggests rapid cooling in Middle to Upper Eocene time, followed by a period of slower cooling after ~30 Ma (Lower Oligocene). The cooling history can also be evaluated by comparison of the ZFT and AFT ages from the same samples. Two samples from the northern and two from the southern part of the study area were analysed for both ZFT and AFT thermochronology. In all four samples, the ZFT ages (50-45 Ma) and the AFT ages (38-32 Ma) are clustered, and therefore show the same cooling history between the ZFT and AFT closure temperatures (~280 to 110 C; Fig. 16B). 48

77 Fig. 15. Topographic profiles with AFT ages from the northern (A) and southern (B) part of the study area (see text for discussion and figure 12 for locations of the profile lines AA and BB ) (Vamvaka et al. 2009). Taking into consideration the uncertainties of the closure temperatures and the error of the ages, an average cooling rate of 13 ± 8 C/myr results for the period between about 50 and 30 Ma. Assuming a reasonable geothermal gradient of C/km, an exhumation rate of 440 ± 70 m/myr results for the mean cooling rate of 13 C/myr. The period from 30 Ma to recent time was characterised by an average cooling rate of 3.2 ± 0.7 C/myr, as can be calculated from the AFT closure temperature and the recent mean surface temperature of ~15 C (Fig. 16B). This results in an exhumation rate of 110 ± 20 m/myr for the mean cooling rate of 3.2 C/myr and the same range of geothermal gradient. 49

78 Fig. 16. A. Central AFT ages (±1ζ) of all samples in relation to elevation. Diagram indicates fast exhumation around 40 Ma and slower exhumation after 30 Ma. B. Central AFT and ZFT ages plotted against closure temperatures (110 C for AFT and 280 C for ZFT); shaded area outlines general trend of cooling rate indicated by the four pairs of FT ages. Fast cooling is indicated from 50 to 30 Ma, followed by slow cooling up to present (Vamvaka et al. 2009). 50

79 Four samples with different AFT ages were used for thermal modelling. Although the samples were derived from different structural levels, i.e., they cooled through the temperature range of the APAZ at different times, they yield rather similar thermal histories (Fig. 17), which are also in agreement with those described above. To get a time constraint for the models at higher temperature, an age of 40 to 60 Ma at 200 C is taken for all model runs. This is based on the ZFT ages from the study area and the western Pelagonian microcontinent in general, which are all around 50 to 45 Ma (Figs 12, 16B). The wider age range for the model runs is due to the uncertainties for the closure temperature. Thermal modelling shows that the samples experienced rapid cooling in the Eocene, followed by much slower cooling in the Oligocene and Miocene (Fig. 17). The broad track-length distribution and the thermal model of sample AV-23 indicate a reheating or at least a prolonged stay in the APAZ between ~30 to 10 Ma (Fig. 17D). This coincides with extension and crustal thinning in the area, which led to the formation of the metamorphic dome in the Olympos region to the east of the study area. The latter sample also indicates an increased cooling rate around 10 Ma. This is indicated in the part of the t-t path which is out of the reliable limits of the modelling results, and thus normally considered as not of great significance, but however it seems realistic since it refers to the period from 10Ma to present, and therefore fast cooling is necessary for the sample to reach the present surface and ground temperature. 51

80 Fig. 17. Graphical presentation of the results of thermal modelling of AFT data for Pelagonian basement samples based on AFT ages and track-length distributions. Modelling was carried out with the HeFTy program of Ketcham (2005). See text ( Sampling and Analytical Procedure and Interpretation and Discussion: Pelagonian microcontinent in the study area ) for further details (Vamvaka et al. 2009). 52

81 Sediments of the Mesohellenic Trough Two main AFT age populations can be recognised in the clastic sediments of the MHT, one around ~40 Ma (Eocene) and the other around ~65 Ma (Paleocene) (Fig. 18, Table 3). The youngest population appears in all samples and is well defined by both the PopShare and BinomFit programs (see above). The four samples with high ρ 2 probability (i.e. AV-81, AV-50A, AV-52 and AV-107), show the same range of AFT single-grain ages and similar distributions to the other samples (Table 3). Therefore, deconvolution of the data was also performed, inasmuch as the ρ 2 probability may result due to overlapping of the different AFT age populations. Statistic results were good and similar to those of the other samples. In this case, the younger AFT age group for the samples AV-52 and AV-107 appears to be close to their central ages; for this reason, the central ages are accepted and used in the further interpretation. On contrast, the distinction of different AFT age-populations for samples AV-81 and AV-50A revealed similar ages to those of samples AV-80 and AV-52, which belong respectively to the same formations (i.e. Pentalophos and Krania Formations) and thus are characterised by similar sedimentation ages. Therefore, the deconvolution of those data was accepted and is further discussed below. Upper Cretaceous ages exhibit a diffuse distribution rather than clusters. This means that in the source area of the sediments, a broad distribution of the Upper Cretaceous to Lower Tertiary AFT ages existed, similar on one hand to what is presently found in the Pelagonian microcontinent in the southern F.Y.R.O.M., but on the other hand extended to higher ages. We therefore refrain from a separation of older age populations (Fig. 18). The Upper Cretaceous to Paleocene ages indicate a source terrain in that part of the Pelagonian microcontinent or in the adjacent Axios Zone, which experienced Eohellenic metamorphism, whereas the Eocene tectono-thermal overprint was so weak that the AFT ages were not or only partly reset. As inferred above, it is also possible that the areas at the eastern Pelagonian microcontinent and the Axios zone, may have experienced a little earlier the main compressional event as this migrated to the west through time. We presume that such areas with Upper Cretaceous to Paleocene ages existed to the east of the Mesohellenic Trough. There, the Pelagonian microcontinent was also overlain by the Jurassic ophiolites, the remnants 53

82 54

83 Fig. 18. Apatite single-grain ages presented in radial plot and probability density plot for the sediment samples from the MHT. The two-coloured (red to orange) bar indicates sedimentation age. For three samples with high ρ 2 probability (AV-50A, Av-52 and AV-107), only one AFT age population is shown. Diagrams with more than one age population are calculated with program PopShare (Dunkl and Székely, 2003). The young age populations showing reheating in samples AV-80 and AV-81 are shown on the corresponding probability density plot. 55

84 56

85 of which are still present as tectonic klippes. In the Olympos area to the east of our study area, however, all possible older ages were extinguished by updoming and extension in the Miocene during the formation of the Olympos Window (see above), so that the older information had been erased. The Middle to Upper Eocene AFT age group around 40 Ma clearly reflects cooling in the source area in response to the Lower to Middle Eocene orogenic event, which heated the according source rocks to temperatures above the AFT closure temperature. According to the AFT age pattern in the southern Republic of Macedonia, the source rocks for these younger apatite crystals are supposed to have been positioned to the west of the source area of the older apatites. That means that the Eocene apatite crystals were derived from a source area along the border of the MHT, where Eocene AFT age resetting and rapid cooling could be demonstrated to have occurred in the basement samples described above. It is interesting that sample Ept-2 from the lower strata of Pentalophos Formation (Upper Oligocene), the only countable sample between those collected from the western part of the MHT, show the same AFT age populations as the other sediment samples which are close to the eastern margin of the trough. Although no other data from the western part were available, nor from the western basement boundary, however the age populations of this sample indicate that this western part of the trough was probably fed by the Pelagonian basement at least since Upper Oligocene times. The lack of sufficient amount of apatite crystals in the other older samples of the western part, belonging to lower stratigraphical levels and of course locating closer to the western MHT margin (see Fig. 11), is an evidence that they derivied from the western ophiolitic boundary rocks; besides, the pebbles of the conglomeratic components of those formations are mainly ophiolitic. Before examining the contribution of the results from the oldest sediment samples (AV-80 and AV-81, Krania Fm., Middle to Upper Eocene), first it is important to discuss the results from sample AV-81. As described above, two different mounts were analysed from this sample, containing different size of apatite crystals (i.e. AV- 81-s, AV-81-b; see results and Table 3). Both mounts passed the ρ 2 probability test. Even though the analysed grains from mount AV-81-s were few, they pointed one AFT age-population of ~40Ma. The mount AV-81-b indicated a central AFT age of ~46 Ma, while the probability density plot for the AFT single-grain ages and also the 57

86 deconvolution of the data show a group of ~40 Ma, clearly distinguished by the older single-grain ages. The mean of the Dpar values from the first mount is 2.7 κm, again different from the average of 3.4 κm of the second mount. All these argue toward the existence of two AFT age-groups, of ~40 and ~60 Ma, in spite of the high ρ 2 probability, probably due to the overlapping of the single-grain ages. The integration of the results from both mounts gave a probability density plot similar to that of AV- 81-b, where the ~40 Ma group is easily distinguished. Comparing the two oldest sediment samples (AV-80 and AV-81), a great similarity of the AFT age groups is observed. The depositional age of the two samples (i.e., Middle to Upper Eocene) is very close to their Eocene AFT age component (Fig. 19). The short lag time between AFT age and sedimentation would confirm the existence of a rapid exhumation pulse in the Eocene as an immediate response to the Neo- Hellenic Eocene orogenic event, which is in line with the results from the FT survey from the adjacent Pelagonian microcontinent. However, in addition to the Eocene age component, both samples show a narrow but distinct cluster of single-grain ages around 30 Ma. These ages show that the samples experienced slight thermal overprint. In the case of an adequate heating to reset AFT ages, one would expect to meet only one AFT age group, younger than the sedimentation age, which is not the picture we get from the results. On first sight, slight reheating in the Lower Oligocene, around 30 Ma, sounds a reasonable scenario, since the Krania Formation experienced compressional deformation, which is not recorded in the unconformably overlying formations. This compressional event should have occurred between ~35 and 30 Ma, so that the young age group of the Eocene sediment samples is excused. Alternatively, a general shift of all AFT single-grain ages to younger ones may have occurred during reheating of those sediments; under these circumstances it would be impossible to presume safely any time of this event. The Dpar values of the samples show a slight difference between grains with different ages (higher Dpar values to higher single-grain ages). However, this difference is not really clear, but very loose, especially for sample AV-80, fact that argues for the partial annealing and shifting of all single-grain ages. This is also backed by the overall AFT age pattern of the sample. In conclusion, this means that no particular time can be deduced for the reheating event. A likely scenario would be 58

87 Fig. 19. Youngest AFT age population (squares) plotted against sedimentation age from the MHT samples. Diagonal lines give lag times in 10 Myr steps. The sample Ept-2 from the north-western part of the trough is not portrayed here since it belongs to a different statigraphic sequence. The possible AFT ages caused by reheating, younger than the sedimentation age of the oldest samples (Krania Fm.) are shown with a triangle symbol; the post-depositional age components from the younger group is presented with a square symbol as for all the other samples. Cycle symbol stands for the samples which display a high chisquare probability, indicating a single AFT population. the Lower to Middle Miocene extensional event, which caused updoming in the Olympos-Ossa region further east. A general shift of the ages does not obviate a short lag time in that sample, since the lag time increases with decreasing depositional ages. The samples from the Lower Oligocene Eptachori Formation and the younger sample from Upper Oligocene-Lower Miocene Pentalophos Formation (i.e., AV-50A, close to the base of Meteora conglomerates) show an increasing lag 59

88 time from about 12 to 22 myr (Fig. 19). No thermal overprint is recorded by the single-grain data. The lag time indicates slow to very slow exhumation in the source area in the order of 0.3 to 0.1 km/myr. In contrast, the directly following decreasing lag time of sample AV-52 which corresponds to the higher levels of Pentalophos Formation, provide an evidence for the extension during the period around the Oligocene-Miocene boundary. Normal faulting may have enhanced exhumation rates, erosion, and feeding of the adjacent MHT with sediments. Taking under consideration the conglomeratic structure of Pentalophos Formation in the particular area, it is rational to surmise a steep part of Pelagonian basement, proximal to the MHT, which started to feed the trough with big amounts of sediments. Denudation, strengthened by extensional tectonic activity, supplied the basin with rocks from higher to deeper structural levels. Increasing lag times in the younger sediments, deposited after 20 Ma, show that the denudation paroxysm took place only around the beginning of Miocene. A very slow exhumation in the source area is indicated from end of Lower to Middle Miocene times. AFT ages around 35 to 40 Ma are still found today in the Pelagonian basement bounding the MHT (see above; Fig. 12). The change from rapid exhumation in the source area of the sediments in the Eocene, and lower exhumation from the Oligocene and onwards, as recorded in the FT ages of detrital apatites, is consistent with the conclusion drawn from the FT survey in the Pelagonain microcontinent along the eastern border of the MHT. In the remote hinterland, in contrast, source areas yielding Upper Cretaceous to Paleocene AFT ages existed throughout the sedimentation period in the MHT. This suggests that exhumation there must have been very slow. The doming event in the Olympos area, where Lower to Middle Miocene AFT ages are present (see above; Most, 2003), is not recorded in the sediments of the MHT, because the youngest sediments that could be sampled in the study area have about the same age as recorded by the AFT data in the Olympos region. 60

89 5.5. Conclusions on the FT analysis The Pelagonian microcontinent is adjacent to the Mesohellenic Trough (MHT) and also constitutes the former source area for Eocene to Miocene clastic sediments, at least of the eastern part of the trough, as signified by the gneissic pebbles of the conglomeratic strata. Apatite and Zircon FT data from the crystalline basement in the western Pelagonian microcontinent reveal significant information on the lowtemperature history of the area (i.e. of the upper crustal levels) for the recent geological time after the Upper Jurassic to Lower Cretaceous orogenic event. AFT data from detrital apatite from the MHT sedimentary formations at the south-eastern part of the trough confirm the Pelagonian microcontinent as the source of the detrital material. They also attest conclusions on the recent exhumation history of the Pelagonian microcontinent; moreover, they provide information on the provenance of the MHT sediments through time and the AFT age terrains in the paleo-source area. The FT data from the basement and the sediments are consistent, and show rapid cooling and exhumation of the crystalline rocks around 50 to 40 Ma (Lower to Middle Eocene) and much slower cooling and exhumation in the Oligocene to Miocene. The Lower to Middle Eocene orogeny caused only partial rejuvenation of the AFT age pattern in the Pelagonian crystalline basement. Heating that clearly overstepped the AFT closure temperature occurred only in the western part of the Pelagonian Zone during this event. Thermal modelling revealed fast cooling through the uppermost crustal levels around Eocene time, which was followed by slow exhumation rates of less than ~0.3 km/myr until about 10 Ma. From combination of ZFT and AFT ages, an exhumation rate of ~0.45 km/myr is calculated for the time between 50 and 30 Ma. In the sediments of the MHT, two main AFT age components around 40 Ma and 65 Ma are recognized. The constant presence of the Eocene age component results in increasing lag times in the younger sediment samples. Short lag time in the Eocene sample supports fast exhumation during the Eocene, as is well documented by the basement FT data. Increasing lag times of all the Oligo-Miocene sediment samples indicate slow exhumation subsequent to the Eocene pulse. FT ages of detrital apatite between ~60 and ~100 Ma probably were derived from a more distant source in the Pelagonian microcontinent or in the Axios Zone. This can be concluded from the FT age pattern in the relevant units in the southern Republic of Macedonia, where a clear AFT age gradient from higher ages in the east 61

90 to lower ages in the west is documented (Most et al. 2001, Most 2003). An Eocene thermal overprint was weak in that area. Apatites displaying higher FT ages may also have been derived from units on top of the actual erosion level in a more proximal position. A prolonged stay of the Pelagonian basement rocks at temperatures of ~100 C is demonstrated by the thermal models based on the AFT data to have occurred between ~30 to 10 Ma. This indicates a stagnation period or may be associated with an uplift of isotherms during crustal extension in response to the Eocene orogenic event. Thermal modelling also reveals final enhanced cooling in the last ~10 my. This is probably related to the termination of the extensional period and drop of the isotherms to normal levels. As every basin, the MHT was supplied with sediments by the surrounding basement rocks (Pindos rocks in the west and Pelagonian rocks in the east), depending on the location where sedimentation mainly took place each period of the trough s evolution, the paleo-morphology, and the transport (river) systems. AFT data shows that the Pelagonian microcontinent supplied with sediments the proximal regions (such as Meteora Nea Zoi region), whilst the sediments at the broader and deeper basin to the west (north in case of Meteora region; see also bathymetry map in figure 4) were probably derived both from the basement rocks, and the reworked material and fine components of the younger deposits. Sample Ept-2 (Pentalophos Fm.) shows that even the western part of the MHT received detrital material from the Pelagonian microcontinent since Upper Oligocene. On the contrary, the ophiolitic basement rocks of the western MHT boundary fed the adjacent western part of the trough almost exclusively during Middle Eocene-Lower Oligocene, as intimated by the deficiency of apatite crystals in the samples coming from this part of the trough, from Krania and Eptachori Fms. This is backed by the locations where conglomeratic deposits occur at this area, which are consisted mainly of ophiolitic pebbles. Summarizing, the Krania and Eptachori formations at the western margin were supplied with sediments from the western bounding basement rocks (i.e. ophiolites, Cretaceous limestones and flysch reworked sediments). Pentalophos Fm. received from both Pindos and Pelagonian rocks, and also reworked material inevitably from the older Eptachori Fm.. The eastern located Tsotyli Fm., as well as all the older formations at the eastern part of the trough (i.e. since Eocene time), was fed by the eastern boundary basement rocks (i.e. Pelagonian microcontinent rocks and the thrusted ophiolites). 62

91 6. STRUCTURAL ANALYSIS 6.1. Methodology Works accomplished in the field and in the office, are reviewed below: - Observations on the tectonic structures, the geomorphology and the stratigraphic relationships have been made. Planar and linear elements have been measured and cross-sections were sketched in the field. Topographics maps of 1: scale (from Geographical Military Service) and GPS were used for the orientation, positioning and marking the observing features, while geological map sheets of IGME (Istitute of Geological and Mineral Research; scale:1:50.000) covering the broader MHT region were used as guides for the geology of the study area. - Study relating to MHT bibliography. - Study about basin analysis and classification of basins in different types according to their geodynamic origin, their development mechanisms and the type of sediments. - Dynamic and kinematic analysis, which comprises: (i) the distinction of different tectonic events; (ii) the construction of Schmidt diagrams for the assessment of the main stresses orientations and therefore the stress regime dominating in each tectonic event. For the paleostress analysis I used both the StereoNett (version 2.46; Johannes Duyster, Ruhr University, Bochum, 2000) and the MyFault programs (version 1.03, Pangaea Scientific, ). The StereoNett program applies P-T method (after Turner 1953), while the MyFault program appears advantageous, and thus preferred, as it provides results by different methods and more detailed information. Firstly, the simple-shear tensor average method was used which is similar to the method developed for analysing calcite twins in deformed rocks (Turner, 1953; Spang, 1972, 1974). With this method, in order to find the stresses causing slip on a set of faults, a simple shear stress state is assigned at each fault, with the intermediate principal stress lying in the fault plane perpendicular to the slip direction. The local stress tensors are then averaged together to give an estimate of the regional stress tensor. The angle between 63

92 the maximum principal stress and the fault plane is varied between 0 and 45 to search for the minimum deviation between the faults in the set. The results of this method were compared with those of the minimized nonslip shear stress method which is based on Angelier method (1979, 1984). The second method takes under consideration that, for a uniform regional state of stress, the direction of slip will depend upon the orientation of the fault and local factors such as frictional anisotropy, and therefore the actual slip direction may not coincide with the maximum resolved shear stress. To estimate the regional stresses, Angelier (1984) examined a number of relations expressing the deviations between the maximum resolved shear stress on the fault plane and the actual slip direction, all of which lead to a non-linear minimization problem. - In the end, transverse cross-sections of the MHT were designed, as well as schematic cross sections that show the different stages of the evolution of the trough. 64

93 6.2. Geometry and kinematics of deformation Field Observations Structures of Krania Formation at the Krania-Spileo villages area (northern Krania sub-basin) Clear compressional structures are evident in the Eocene strata of the Krania Formation. Asymmetric folds follow the shape of the norhtern Krania sub-basin (at the western part of the MHT, inferred as Gulf of Krania or Krania basin after Brunn 1956 and Wilson 1993, respectively); the b-axis of the folds has a main NNW-SSE orientation at the centre of the basin, and is curved towards WSW-ENE at the northern margin of the sub-basin (see Fig. 20). Reverse faults mainly dip to SW with SW-plunging dip-slip striations. In places, reverse faults dip also to NE with NEplunging dip-slip striations accordingly. Kinematic indicators reveal that the main sense of movement is towards the NE to ENE (Figs 20 to 22; Pict. 1). The facies of Krania Formation, coarse (mainly ophiolitic) breccias and conglomeratic fan-deltas turning to turbiditic deposits, reveal a steep profile and rapid subsidence. The ophiolites and the Krania strata are almost concordant, vertical or dipping at a high angle (ca. 80 ) to the east or west; their contact is highly tectonized. The northern and southern margins of the northern Krania sub-basin are bounded by strike-slip faults, as previously documented (Papanikolaou et al. 1988; Doutsos et al. 1994; Ferrière et al. 2004). The folding and reverse faulting have been recognised in previous studies (e.g., Wilson 1993, Doutsos et al. 1994, Vamvaka et al. 2006), while the contact between the Krania Formation and Mesozoic ophiolites beneath, was described as a thrust (Wilson 1993; Doutsos et al. 1994) or a sub-vertical flexure, locally reverse and faulted (Ferrière et al. 2004). The structures observed at the area of the northern Krania sub-basin, especially at the contact with the bounding ophiolitic rocks, are described analytically below. 65

94 Fig. 20. Northern Krania sub-basin, where the strike and dip of Krania strata, and the observed structures are shown. The lighter yellow colour represents the continuation of Eocene Krania formation; it is unknown which exactly is its extend underneath the younger deposits. Bedding of Krania Fm. and b-axes of the folds are illustrated in diagrams. Numbers indicate observation locations discussed in the text. MIKR: village Mikrolivado; MON: village Monachitis. Red lines indicate cross-sections of figures 23 to

95 Fig. 21. Photograph and sketch of reverse fault cutting Krania formation strata (northern Krania sub-basin, south of Krania village and number 1 as shown on the map of figure 20). Fig. 22. Reverse faults deforming Krania formation strata at the western part of MHT (near Krania village, next to the photograph of figure 21). 67

96 Pict. 1. Folding with NW-SE orientation of b-axis and reverse faulting parallel to the axial plane of the fold, at the middle part of the Krania sub-basin area (east of Krania village and number 1 as shown on the map of figure 20). At Krania village, the Eocene strata adjacent to the contact with ophiolites is described by its steep dipping angle and a gentle waving along strike (number 1 in figure 20). The tectonic contact between Eocene strata and the ophiolitic rocks is ca. 90 and has NW-SE orientation. Strike-slip striations were found in places, both on the ophiolites and on calcites grown parallel to the bedding of adjacent to the contact marls and sandstones, which reveal a shear character in the contact zone. This strike-slip movement is probably related to the waving form of the bedding (Figs 20 to 23, Picts 1 to 7). On the other hand, the high inclination of the deformed Krania Fm. strata signifies clearly that the formation experienced compressional tectonics. Dip-slip striations were observed in places on the ophiolite fault surfaces at the contact with Krania Formation, associated with normal movement and resulting in the lowering of the eastern side. This normal movement is considered younger, developed at a later stage due to reactivation of the older fault planes. 68

97 Pict. 2. The contact of Krania Fm. with the ophiolites bounding the MHT to the west (Krania village; number 1 in figure 20). Picts 3, 4. Almost vertical Krania Fm. strata (i.e., marls and sandstones), in contact with the basal ophiolitic conglomerates (number 1 in figure 20). 69

98 Pict. 5. Reverse fault cutting the Krania Fm. strata (profile); striation with sinistral strike-slip component (number 1 in figure 20). Pict. 6 (left). Dextral strike-slip fault on the basal ophiolitic conglomerates of Krania Fm. Pict. 7 (right). Dextral strike-slip fault cutting the Krania Formation adjacent to the contact with the basement; calcites are grown parallel to the fault surface (number 1 in figure 20). 70

99 Fig. 23. Cross section at the contact of ophiolites with Krania Fm. strata, at Krania village (number 1 in figure 20). White arrows are used to show the younger normal movement on the older faulted surfaces. Westwards of Krania village, on the way to village Milia (number 2 in figure 20), the contact of the ophiolites basement rocks and the ophiolitic conglomerates of the base of Krania Fm. is crossed. Krania strata appear almost vertical once more, while WNW-ESE to NW-SE trending faults at the contact of basement rocks and Krania Fm denote an oblique reverse movement (Fig. 20). Sinistral faults with a normal component, striking NE-SW, also occur at the same location. Northern of Krania village, northwest of the village Mikrolivado, the contact of ophiolites with the Krania Fm is also described by an oblique reverse fault. An ophiolitic block is observed on top of Krania Formation, very close to the contact of the sediments to the basement (number 3 in Fig. 20, Pict. 8, Fig. 24). This is caused by a big gravitational sliding of the ophiolitic mass onto the sediments. The sediments exactly under the ophiolitic block appear with a ca. 50 angle, dipping to the east. Sinistral strike-slip faults of NE-SW orientation were observed at the southern margin of Krania sub-basin, along the contact between ophiolitic rocks and Eocene strata (Fig. 20; number 4). Along the main northern boundary of the Krania sub-basin (number 5 in figure 20), a dextral strike-slip fault with almost E-W strike is documented by Papanikolaou et al. (1988). Large limestone and ophiolite olistoliths occur all along the southern part of this fault. Ferrière et al. (2004) report the fault of the northern margin as Monachiti-Trikomo structure (MTS), because of the villages names in the area. The specific scientists do not declare themselves towards a dextral or sinistral movement, but support that this strike-slip fault was active at least 71

100 since Upper Eocene, because (i) same age sediments are different on both sides of the MTS, (ii) the middle Krania sandstone rest unconformably on the northern side of this structure, which is the source of the olistolithic channel-fills at its southern side, and (iii) the Oligocene Eptachori strata truncate the uppermost, subvertical upper Eocene turbiditic strata of Krania Fm. onto the southern flank of the MTS. A lateral-normal fault of almost E-W orientation with dextral strike-slip component has been observed also at the northern-most part of Krania sub-basin (Fig. 20; number 6), nearby the contact with the Eocene basal conglomerates. Pict. 8. An ophiolitic block placed upon the Krania Fm. strata, due to sliding (number 3, Fig. 20). Fig. 24. Cross section at the contact of ophiolites with Krania Fm. strata, close the village Mikrolivado, where an ophiolitic block is emplaced on top of Krania Fm. (number 3, Fig. 20). Same legend as in figure

101 In that northern-most small area, west of Spileo village, Krania Formation is enclosed by ophiolites and the overlying Cretaceous limestones which seem to form a positive flower-shaped structure (Fig.20; number 6). In most places, as at the northern part, the strata lies on the basement rocks with a steep angle, and dipping at the same orientation as the basement rocks. On the contrary, at the middle part of this area (number 7 in Fig. 20, and Fig. 25), the ophiolites seem to thrust with a SW- NE orientation over the Eocene strata with an angle of ~80. Sinistral faults of SW- NE orientation, dipping with a high angle towards NW and elsewhere towards SE, with reverse and normal component respectively, were observed cutting the ophiolites at this location; reverse faults with a dextral component striking WNW-ESE, also occur. Strata at the specific location dip to NE with a high angle (~65-70 ) and cut by a fault striking the same orientation (i.e. NW-SE) and dipping to SW. No striation could be observed, but the continuation of a sandstone bed on either sides of the fault plane denotes a reverse movement (Fig. 25). Fig. 25. Cross section at the southern contact of ophiolites and Krania Fm., in the small enclosed area at the west of village Spileo (number 7 in Fig. 20). The ophiolites thrust against the Eocene marls along WSW-ENE orientation, while several sinistral faults of this strike observed at the same location. Moving eastwards of this location, towards Trikomo village area, more sinistral faults of SW-NE orientation were observed at the ophiolitic rocks and also at the lower ophiolitic conglomerates of Krania Fm. At the NE-ern limit of Krania sub-basin, another fault with almost the NE-SW strike is documented on map sheet Panagia of IGME (1980). It occurs at the contact between Krania and Eptachori Fms, which indicates that it was active after the deposition of Krania Fm. (i.e. after Eocene times; Fig. 26). 73

102 At the southern side of this fault (i.e. south of the village Trikomo ), and also near the contact with ophiolites, Krania strata is observed again with a waving along bedding, apparently due to tectonic activity (Fig. 20, number 8; Fig. 26; Pict. 9). Younger breccias and conglomerates with reddish colour, recorded as Upper Eocene fan deltas (Zelilidis et al. 1997, 2002), lie unconformably on top of the upper Krania deformed fine sediments, with a gentle angle towards east (Fig. 26). It is noteworthy that at the particular location (south of Trikomo village), the Eocene strata dips towards SSE or SSW, curved in relation to its general NW-SE strike (see Fig. 20). Fig. 26. Cross section at Trikomo village area (number 8 in figure 20). Faults account for the relatively uplifted Krania Formation in comparison to the ovelying Eptachori Fm. The Uppermost Eocene breccia of Eptachori Fm. lie unconformably on top of the highly deformed Krania strata. Pict. 9. Almost vertical Krania Fm. strata dipping SSW, with a small waving along strike south of Trikomo village (number 8 in figure 20). 74

103 Structures of Krania Formation at the Vassiliki Nea Zoi villages area (Southern Krania basin) The Krania Fm. at the south-eastern part of the Mesohellenic trough shares similarities with the same age formation at the western part of the trough (i.e. northern Krania sub-basin ). The Eocene strata at the southern Krania sub-basin appear to dip towards SW and elsewhere towards NE, in places with steep angles (~70 ), forming a folding with b-axis trending NW-SE (Fig. 27). Although no reverse faults were observed, the folding and in places the almost vertical bedding reveal that the area was affected by compression during Eocene times. The base of the Eocene formation at the eastern side of this southern Krania subbasin is consisted by a thin layer of carbonate basal conglomerates, derived apparently from the Triassic carbonates of the Pelagonian basement (Nea Zoi village), on which is lying trangressively (Fig. 27, Picts 10 and 11). These first deposits designate the beginning of deposition in a coastal environment at the eastern side of the sub-basin, after sea transgression (area of Nea Zoi village). Subsidence and sedimentation rates should be high as revealed by the FT analysis and also the rapid change of facies from conglomeratic to marly nummulitic limestones (Picts 12, 13) and to turbitidic marls. Fig. 27. Southern Krania sub-basin, where the strike and dip of Krania strata, and the faulted western margin of the sub-basin are shown. The lighter yellow colour areas represent the possible continuation of the Eocene deposits. Bedding of Krania Fm. and b-axes of the folds are illustrated in diagrams. THEOP: Theopetra village. 75

104 Picts 10, 11. Basal carbonate conglomerates at the base of Krania Fm. at the eastern margin of the south Krania sub-basin, including also bigger angular carbonate fragments fallen from the Pelagonian basement rocks (area of Nea Zoi village). Picts 12, 13. Nummulitic limestones and marls of the Eocene Krania Fm., at Nea Zoi (left) and Vassiliki (right) villages areas, forming the first marly-sandstone beds over the basal conglomerates. The western part of the sub-basin is bounded by ophiolites and Cretaceous limestones; at Vassiliki village area; carbonatic breccias occur over the ophiolites, south-east of Vassiliki village (Fig. 27, Pict. 14). The western contact of the sub-basin, between the basement rocks (Cretaceous limestones and ophiolites) and the Eocene strata is tectonized, similar to the northern Krania sub-basin, as evident by steep strike-slip faults observed at the contact of the Cretaceous limestones (and the underlying ophiolites) and the Eocene marls near Theopetra village striking NW-SE (Fig. 27, Pict. 15). Both dextral and sinistral movements could be observed on the 76

105 faulted zone, showing that the same fault was reactivated in later stages of the MHT evolution. As it resembles the western margin of the northern Krania sub-basin, we presume that the Eocene strike-slip faults of NW-SE orientation, bounding the southern Krania sub-basin to the west, have also a dextral component. Pict. 14. Carbonatic breccias overlying the ophiolites at the western part of the south Krania subbasin (south-east of Vassiliki village). Pict. 15 (right). Strike-slip faults at the contact between the Cretaceous limestones and the Eocene marls, at the western margin of Krania sub-basin (near Theopetra village). 77

106 Normal faults observed in the area of Nea Zoi village, and striking NW-SE, cut both the Triassic and the Eocene limestones (Picts 16, 17 Fig. 28), and therefore developed after the deposition of Krania Fm. (i.e. after Eocene). Picts 16, 17. High-angle normal fault of NW-SE orientation, cutting the Triassic limestones and the Eocene strata (near Nea Zoi village). Fig. 28. Cross section at Nea Zoi village area (eastern part of the southern Krania subbasin). Normal faults cut both the Triassic limestones and the Eocene nummulitic limestones, signifying its activity after Eocene times. 78

107 Structures of Eptachori and Pentalophos Formations area Eptachori Formation is generally described by strata dipping to the east, with a steep angle of ~60-70 near its contact with the basement rocks. The first fan-deltas deposits of late Eocene-beginning of Oligocene times (as reported by Zelilidis et al. 1997) lie unconformably on Krania Formation, while its contact with the basement is largely defined by a steeply dipping fault between the basement rocks and Eptachori Fm., forming an impressive morphology. This fault manifests a dextral strike-slip movement on NW-SE orientation, with both normal and reverse dip-slip components, as well observed at many places along its development between the basement rocks and Eptachori Fm. Dextral strike-slip faults with clear striation and in places with nicely developed shear-bands were observed (see Fig. 29) at the northern-most area at Grammos Mountain (Picts 18, 19; number 1 in Fig. 29), and going southwards near the villages Dotsiko (Picts 20 to 23; number 2 in Fig. 29), Mesolouri (Picts 24, 25; number 3 in Fig. 29), Alatopetra (Pict. 26; number 4 in Fig. 29), Spileo (number 5 in Fig. 29), and Chani Mourgani (Picts 27, 28; number 6 in Fig. 29). At the southern part of the MHT, where the Triassic to Jurassic limestones and superimposed ophiolites compose Koziakas Mountain that borders the western side of Mesohellenic trough, strike-slip faults are also observed (number 7 in Fig. 29). Photographs of the marginal fault along the western boundary and the striations observed on the fault planes are displayed in the next pages. The dextral strike-slip zone along the western boundary of MHT trough appears a small divergence of its strike from the main NW-SE orientation, which is usual for strike-slip faults (Fig. 29). A second dip-slip striation overprinting the strike-slip one is observed on many fault surfaces. Dip-slip striation on cataclastites which are superimposed on strike-slip striations, show that strike-slip preceded normal faulting. 79

108 Fig 29. Eptachori and Pentalophos formations; the dashed-lined area indicates the possible extension of the basin and depositional area during Oligocene - Lower Miocene. The numbers indicate the locations discussed in the text. Tectonic contacts and relative movements are illustrated with thicker black lines and arrows; 1: Strike-slip faults; 2: reverse faults/ thrusts; 3: normal faults. Grey colour is used for the structures of Krania Fm. area. Red line at number 8 indicate crosssection of figure

109 Pict. 18: The contact of basement rocks forming a high relief with the Eptachori Fm strata dipping to NE (Grammos mountain). Pict. 19: NW-SE trending dextral strikeslip fault with a small normal component, on the Cretaceous limestones of the location Vertenik at mountain Grammos. Pict. 20: NW-SE dextral and NE-SW sinistral strike-slip faults cutting the ophiolites, next to the contact with the Oligocene Eptachori formation, at a stone-pit at the area of Dotsiko village (north from Filippei village; number 2 in Fig. 29). 81

110 Picts 21, 22 (Top). Dextral NW-SE trending strike-slip fault with normal component at the ophiolites near the contact with the Eptachori Fm. (profile). Striations have been observed on the faulted surface (Dotsiko village area, same location as picture 20). Pict. 23A,B. Sinistral NE-SW trending strike-slip fault with normal component at the same location (plan view; xz-plane). 82

111 Pict. 24. The tectonic contact of the Cretaceous limestones with the Lower Oligocene Eptachori Fm. strata of the MHT, near the village Mesolouri (north-east of Filippei village; number 3 in Fig. 29). Pict. 25. The Lower Oligocene strata of Eptachori Formation near the contact with the basement rocks, dipping to the east with a high-angle angle (~60-70 ) (near Dotsiko village; number 2 in Fig. 29). 83

112 Pict. 26. The marginal strike-slip fault of the western basin boundary (Alatopetra village area; number 4 in Fig. 29). The arrows show the two movements occurred in different periods (based on striations found on the fault plane; see text for explanation). Pict. 27, 28. The marginal strike-slip fault of the western basin boundary at Chani Mourgani village area (south-western part of the study area; number 6 in Fig. 29). The nearly horizontal striation on the fault plane on ophiolites reveals a dextral strike-slip character, very close to the contact with the Eptachori Formation. 84

113 The Cretaceous carbonates at village Spileo rise up in a narrow area, subsiding steeply on both sides, west and east, where Krania and Eptachori formations are respectively deposited (number 5 in Fig. 29; Fig. 30; Picts 29, 30). This was previously described as a pop-up structure due to east-verging thrusts (Doutsos et al. 1994), while another interpretation ascribe the development of this structure, as well as the whole western MHT boundary, to basement flexure, normally faulted at its eastern side (Ferrière et al. 2004). However, no reverse movement has been detected on the marginal fault surfaces. On opposite, clear strike-slip striations were found on the almost vertical fault plane at the eastern side (Pict. 31). In addition, the numerous parallel faults of NW-SE trend, and the very steep relief common result of strike-slip faults, argue for the formation of this peculiar morphology due to strike-slip activity, forming a positive flower structure. The shape of the faulted carbonates is also impressing, pointing out a dextral lateral movement (Fig. 30). A second dip-slip striation observed also at those surfaces and showing normal movement is posterior to the strike-slip one, as described earlier. Fig. 30. Spileo village area where dextral strike-slip faults trending NW-SE occur between the Cretaceous limestones and the Lower Oligocene strata of Eptachori Fm. of the MHT. (Geology based on the map sheet Grevena of IGME, 1972). The tectonic contacts and faults are shown with thicker lines. Oph: ophiolites, Lim-Cr: Cretaceous limestones, Fl: Pindos flysch, Kr: Krania Fm., Ept: Eptachori Fm. [Uppermost Eocene (Ept-Fan deltas) Lower Oligocene (Ept)], Pent : Pentalophos Fm., Qt: Quaternaty deposits. 85

114 Pict. 29, 30. Faulted Cretaceous limestones of Spileo village area, forming a steep morphology; they bound the Lower Oligocene strata of Eptachori Fm to the east (first and second photos) and the M.-Up. Eocene strata of Krania Fm to west (left edge of second photo). Pict. 31. The nearly horizontal striation on the dextral strike-slip fault plane on the Cretaceous limestones at Spileo village area, close to the contact with the Eptachori Formation. 86

115 Pentalophos Formation lies conformably on top of Eptachori Fm., dipping also to NE (Brunn, 1956, Savoyat et al. 1972), and appearing similar structures; this denotes a rather stable tectonic environment during the deposition of those two Oligocene Lower Miocene formations. Slumps in turbiditic sediments are very common in Eptachori Formation (Picts 32, 33). Some reverse faults occur in places, striking NW- SE to W-E, as observed mainly in Eptachori and also lower Pentalophos Formations (Fig. 29, Picts 34, 35). Open and gentle folding is also present in few locations (e.g., near Morfi village), with an almost horizontal b-axis plunging NW. The occurrence of reverse faults and gentle folding reveals a compressional influence during the deposition of the two Oligocene Lower Miocene formations. Pict.s 32, 33. Typical slumps in Eptachori Fm, (northern part of the MHT, east of Mountain Grammos), denoting rapid turbiditic sedimentation at steep slopes developed due tectonic activity. Pict. 34, 35. Reverse faults trending NW-SE in Eptachori Formation, near village Prosvoro (north-east from Filippei village; between numbers 3 and 4 in Fig. 29). 87

116 Dextral strike-slip faults with slightly different orientation (NNW SSE) and small reverse dip-slip components also occur towards the centre of the southern part of the MHT (e.g. Theotokos village area; Fig. 29, number 8), where an uplifted flowershaped structure is formed that exposes basement rocks (i.e. ophiolites and Cretaceous limestones) and the Eptachori Formation (Fig. 31). West of this structure the strata dip to WSW, while at the eastern side the strata dip to ENE. The dextral strike-slip faults of this region trend parallel to each other and appear to affect the Eptachori and Pentalophos Formations, whereas the Miocene Tsotyli Formation does not appear to be affected. A positive flower structure is also recognized by Zelilidis et al. (2002) according to their interpretation given to the seismic profile of the specific area showing the strike and dip of strata as well as several unconformities. Fig. 31. Panoramic picture and cross section at Theotokos village area (number 8 in figure 29) where parallel dextral strike-slip faults occur, bounding the Eptachori Formation. 88

117 Parallel strike-slip faults striking NNW-SSE also occur at the ophiolites and Cretaceous limestones protruding south of Kalambaka city (Fig. 29, number 9). This fault zone occurs exactly on a straight line at the continuation of the Theotokos structure towards south, while the between area is covered by Miocene deposits which are not affected by those faults. This could correspond to different segments of the same tectonic zone or reflect even to a great growth of the same NNW-SSE fault, continuing blindly all the way from Theotokos to Kalambaka area under the Miocene sediments, which are not affected by those faults. Dextral strike-slip faults of NNW- SSE orientation have been observed in the ophiolites south-east of Vassiliki village; in this area they define the contact between the ophiolites and the Eptachori Formation to the east (Fig. 29; number 9). More strike- slip faults, both dextral striking NNW-SSE and sinistral striking NE-SW, occur also at the Cretaceous limestones of the region (Theopetra region). Sinistral strike-slip faults, striking NE SW to ENE WSW were often observed at the same locations with the main dextral strike-slip faults, and elsewhere cutting Eptachori and Pentalophos Formations (Fig. 29). Except the main dextral strike-slip fault of NW-SE orientation, developed along the western boundary of MHT, also sinistral strike-slip faults of the same orientation were observed in few places. Such faults occur on the ophiolitic assemblage close to Eptachori village area, adjacent to the contact with the Eptachori formation (number 10 in Fig. 29; Fig. 32); dextral strike-slip faults trending NE-SW also occur as antithetic Riedel faults to the main sinistral zone. The basement rocks at this location consist of dolerites and syn-tectonised Upper Jurassic to Lower Cretaceous limestones (age of limestones is taken from the geological map sheet Pentalophos of IGME, 1960). Besides the strike-slip faults, these rocks are also cut by conjugate reverse faults dipping to NE and SW; the reverse faults are certainly older than the strike-slip ones, as they are clearly cut by them. A little further from this location and very close to the MHT boundary, the Upper Eocene fan delta deposits of Eptachori Fm. (Kontopoulos et al. 1999) change dipping direction several times in a close distance, creating folding with b-axis plunging NE (i.e., sub-horizontal b-axis of ~10 to 60 strike; number 11 in Fig. 29). 89

118 Fig. 32. Photographs and sketches of the syn-tectonised ophiolites and Upper Jurassic-Lower Cretaceous limestones of the western MHT boundary, west of Eptachori village (Fig.29, number 10). Sinistral strike-slip faults of NW-SE orientation is cutting the basement rocks, while the Up. Eocene Eptachori strata lie exactly few metes at the right of the right photograph. Surface weathering and a stream of seasonal flow do not allow obtaining a clear image of the contact. 90

119 In the same area, a lateral reverse fault of the same orientation with important sinistral component was observed in the Eptachori Fm, near Eptachori village area and close to the contact with the basement rocks (Fig. 29, number 11, Picts 36, 37). Sinistral movement along NW-SE orientation was also observed in another location at the ophiolites of the western MHT boundary, near the contact with the Eptachori Formation; at the Mesolouri village area (number 3 in Fig. 29). The picture was not very clear, as both dextral and sinistral strike-slip faults on the same orientation were observed. Faulted surfaces also of NE-SW orientation exhibited elsewhere a normal dextral movement and elsewhere a sinistral reverse movement. The relative age of the two different movements were revealed some distance further by overprinting criteria on faulted surfaces striking NW-SE. Striations dipping gently to east, related to a dextral strike-slip movement, were cut by younger striation dipping also with a low angle to the east, related to a sinistral reverse movement, and therefore the dextral strike-slip fault of NW-SE orientation preceded the sinistral one. Close to the centre of the trough, at a location where Pentalophos Fm. slightly outcrops out of the Quaternary sediments that cover the particular region (east of Grevena city; number 12 in Fig. 29), syn-sedimentary reverse faults trending NE-SW and NW-SE were observed (Fig. 33); the faults of the latter orientation also exhibit an important sinistral component. The synsedimentary activity is revealed from the faults of NE-SW orientation which doesn t intersect everywhere the strata up to all its growth. Furthermore, on two fault planes of NW-SE orientation (Fig. 33), the activation of a second normal along the same surface has been identified. The normal fault is younger than the reverse one as it can be revealed by superimposing calcites growth associated to the two opposite movements. No other reverse fault striking NE-SW has been observed elsewhere, except the location described above. However, it should be mentioned that a couple of faults striking NE-SW are showed in the geological map sheet Grevena of I.G.M.E. (1972), cutting the Pentalophos and Eptachori Fms, and are illustrated as reverse in the cross section of the map; those faults are not continuing to the Miocene Tsotyli Formation. In the end, normal faults striking different orientations have been also observed in Eptachori and Pentalophos Formations. Few examples are shown in the next photographs (Fig. 34 and Picts 38, 39). 91

120 Picts 36, 37. Lateral reverse fault with sinistral component, trending NW-SE, in Eptachori Formation, close to the western basement boundary (north-west of Eptachori village, number 11 in figure 29). Left: the reverse fault cutting the strata dipping to SW; Right: the striation on the fault surface. Fig. 33. Cross section of an outcrop of Pentalophos Fm, in the centre of MHT, cut by reverse faults were observed (number 12 in figure 29). 92

121 Fig. 34. Picture and sketch of the normal faults in Pentalophos Formation, trending from WNW-ESE to ENE-WSW, at the northern part of MHT (south-west of Nestorio village; number 13 in Fig. 29). Picts 38, 39. Normal faults trending approximately E-W, in Pentalophos and Eptachori Formations. Top: two antithetic normal faults, where the vertical displacement of the left fault is more than a meter, as shown on the picture (Pentalophos Formation; number 13 in Fig. 29); south-west of Nestorio village. Bottom: normal fault, south of Trikomo village (Eptachori Formation; number 14 in Fig. 29). 93

122 Structures of Tsotyli Formation area (Eastern MHT margin) The Miocene Tsotyli Formation is extended along the eastern part of MHT. Lowangle normal faults with a small strike-slip component were observed at the eastern boundary of the MHT. These faults occur at the contact between the Tsotyli Formation and the eastern MHT boundary rocks, either the Triassic carbonates (e.g., close to village Paleaokastro ; Fig. 35; Fig. 36, number 1) or the gneissic schists (e.g., close to village Kerasoula ; Fig. 36, number 2) or even the ophiolitic rocks (e.g., area of Pylori village; Pict. 40; Fig. 36, number 3); they show synsedimentary activity but do not affect the younger, Pliocene deposits. Those normal faults exhibit in general a NW SE strike (e.g. Paleokastro, Pylori, Kerasoula village areas) and a southwestward sense of movement. In some places, NW dipping low-angle normal faults show an opposite sense of movement towards NE (e.g., east of Kerasoula village; Fig. 36, number 4; Fig. 37). The low-angle faults have contributed to the subsidence of the eastern part of the basin where the Tsotyli Formation was deposited and probably attributed to the further exhumation of the Pelagonian block. They are not rendered for observation everywhere along the eastern margin of MHT, because in many places they are cut by younger high-angle normal faults or covered by younger post-miocene deposits. Fig. 35. Cross section near Paleokastro village (number 1 in Fig. 36). A low-angle normal fault of NW-SE orientation characterizes the contact between the Miocene Tsotyli conglomerates and the Triassic carbonates of Pelagonian microcontinent. In some distance to the east, the Triassic limestones are also cut by younger high-angle normal faults of the same strike (right fault on this figure; explanation later in the text). 94

123 Fig. 36. Tsotyli Formation, deposited at the eastern part of the MHT during Lower-Middle Miocene, and the low-angle normal faults with small strike-slip component, trending mainly NW-SE along the eastern boundary of MHT between the basement rocks and the Tsotyli Fm. The numbers indicate the locations discussed in the text. Pict. 40. Low-angle shear zone observed at the ophiolites, close to their contact with the Tsotyli Formation (eastern margin of the MHT, close to Pilori village; number 2 in Fig. 36). 95

124 Fig. 37. Photographs and cross-section of the low-angle shear zone trending NE-SW at the gneissic Pelagonian rocks at the contact with the Tsotyli Formation (eastwards of Kerasoula village, number 4 in Fig. 36). Sub-horizontal striation on the fault plane shows a movement towards NE. The low-angle faults develop sub-parallel to the old schistosity of the deformed basement. In the Miocene strata of the Tsotyli Formation, a small number of dip-slip reverse faults striking NW SE, were observed. Dextral strike-slip faults of N-S to NNE-SSW orientation were also observed cutting the Miocene deposits. Sinistral faults of E-W to NW-SE orientation also develop as Riedel to the dextral ones at the same locations. Such strike-slip faults were observed near Nestorio village, where they occur at the Middle Miocene carbonate deposits of Ondria Fm (i.e. Picts 41, 42; Fig. 38, number 1), and near Krystallopigi village, where they cut the Lower-Middle Miocene Tsotyli Fm. conglomerates (Fig. 38, number 2). 96

125 Picts 41, 42. Dextral strike-slip striation on Miocene limestones of Ondria Fm near Nestorio village (number 1 in figure 38). Two other locations exhibiting strike-slip faults cutting Miocene deposits occur at the eastern part of MHT. The first location is near the village Tsotyli (Fig. 38, number 3), cutting the Miocene deposits and the ophiolites outcropping there. The second location is near the villages Paleokastro and Taxiarchis, where a dextral shear zone with NNE-SSW orientation develops at the Miocene ophiolitic conglomerates that compose the base of Tsotyli Fm. (Fig. 38, number 4). The serpentine cobbles of the Tsotyli Fm. basal conglomerates at this location are characterized by silicified rinds, which could have resulted either by their chemical alteration in submarine canyons due to hydrothermal activity or by the impact of sea-water during their deposition on a shallow environment close to sea level that led to the silica concentration. Taking under consideration the second possibility, the sedimentation was taking place (at least at the specific location), in a shallow environment near the contact to the basement rocks in the Aquitanian, when the Tsotyli Fm. started depositing. In the close vicinity to those conglomerates, laterites outcrop, which are nearby covered by the overlying Miocene Tsotyli Fm. coarse sandstone deposits. The laterites intimate a surface exposure of the ophiolites at the atmospheric conditions during Aquitanian or a little earlier, before the deposition of Miocene strata on top of them. The above remarks denote that the subsidence of this eastern area of MHT started around Aquitanian, when the first sediments of Tsotyli Fm were deposited, still in a rather shallow environment. Consequently, this is a proof that the Mesohellenic Trough was not extending that far to the east during Upper Eocene-Oligocene times; it widened towards east through time, especially during Lower Miocene, principally due to the low-angle normal faulting, evident along the eastern boundary of the trough. 97

126 Fig. 38. Faults developed in the whole area of MHT after the Middle Miocene. The youngest MHT formations, Tsotyli and Ondria, deposited during Early-Middle Miocene along the eastern part of the trough is high-lighted with yellow colour. The numbers indicate the locations discussed in the text. The red line south of Theotokos village correspond to the cross-section of figure

127 High-angle normal faults that strike in several different orientations cut the basement rocks, the MHT formations and Plio-Quaternary deposits; they also overprint all previous structures (Fig. 38). Some of these normal faults, generally those oriented E-W up to NE-SW, are believed to be still active (Caputo and Pavlides 1993, Chatzipetros 1998; Chatzipetros et al. 2005; e.g., number 5 in Fig. 38). Normal faults of NW-SE to NNW-SSE orientation dipping to the west are observed between the Triassic carbonates of Pelagonian microcontinent and Pliocene- Quaternary deposits, and cutting all Eocene to Miocene deposits at several places as: north and south of the city Kastoria (e.g., numbers 6 and 7 in Fig. 38; Fig. 39), at the centre of the trough close to Paliouria village (number 8 in Fig. 38), and at the southern part near the village Nea Zoi (Fig. 38, number 9). In many places, the marginal low-angle faults are cut by those high-angle normal faults (Fig. 39). At the western boundary of Mesohellenic trough, normal faults of the same strike are developed in many places, on pre-existed fault surfaces, overprinting the former movements (e.g., number 10 in Fig. 38). Another normal fault of the same orientation and with a small sinistral component is traced east of northern Krania sub-basin, at the contact of the Upper Eocene breccias and the overlying Oligocene marls of Eptachori Fm. (number 11 in Fig. 38). The contact of those coarse and fine Oligocene deposits is followed in big extend by the rivers run in that area, implying the possible existence of a tectonic zone. Fig. 39. Sketch of the NNW-SSE trending high-angle normal faults, cutting the Triassic limestones and the older low-angle normal fault (north from city of Kastoria). 99

128 Approximately at the middle of the trough, west of Taksiarchis village, a normal fault of the same NNW-SSE orientation occurs between an ophiolitic outcrop and the Quaternary deposits (Pict. 43; number 12 in Fig. 38). Further to the south, the contact between the Pentalophos and Tsotyli Formations, along which the Aliakmonas River runs, may be also characterized as a normal fault. Pict. 43. Normal fault, trending NNW-SSE, between the ophiolitic basement rocks that outcrop in the specific area and the Quaternary deposits (west of Taksiarchis village, see Fig. 38 for location). Following the same imaginary line to the south, high-angle normal faults with a NNW SSE orientation dip to the east and west some distance from Theotokos village (Fig. 38, number 13; Fig. 40). The high angle normal faults striking NNW-SSE and dipping to the east occur at the western part of the Eptachori Fm. outcrop, while parallel antithetic normal faults dipping to the west develop at the eastern part of Eptachori Fm outcrop, resulting to the morphologic lowering of this area where Eptachori Fm is exposed (Fig. 40). At this area, the Miocene Tsotyli formation overlie directly on the Eptachori Fm. without the interposition of Pentalophos Fm. The normal faults dipping to the east cut the Eptachori and the overlying Pentalophos Formations, and initially thought to 100

129 account for the direct juxtaposition of the Tsotyli with the Eptachori Formation, caused by the lowering of the eastern region and erosion of the overlying Pentalophos Fm (Vamvaka et al., 2006). However, further investigation revealed this is not the case, as the specific faults (dipping to the east) do not occur at the contact of Eptachori and Tsotyli Formations, but within Eptachori and Pentalophos Fms (see Fig. 40). Fig. 40. Photograph and schematic cross-section of the high angle normal faults of NNW- SSE orientation, observed southern of Theotokos village, which cut the Eptachori and Pentalophos Formations at the western part, and also the Tsotyli Fm. at the eastern part of the figure (number 13 in figure 38). The three Formations are marked on the photograph and the cross-section. 101

130 6.2.2 Tectonic Events Several sets of structures, as described above, record a complex deformational history that is difficult to be unwrapped. Overprinted criteria, stratigraphic data and geometry of kinematics, as well as correlation between the various structures allow us to distinguish six tectonic events which affected the study area. These events took place in semi-ductile to brittle conditions from Middle Eocene to Quaternary time. To assess the stress regime governing each deformational event, we have calculated its stress tensor from a large amount of fault-slip data. Both P-T method after Turner (1953) and Angelier (1979, 1984) methods were used for the paleostress analysis. The two methods showed same or similar results, without exhibiting important differences. The palaeostress analysis diagrams illustrated in the figures were made with the first method. The palaeostress analysis diagrams of both methods and the corresponding stress-inverse solutions are provided in Appendix B. T1 event. Krania Formation experienced considerable NE-SW shortening as demonstrated by folding and reverse faulting. In contrast, the younger formations, Eptachori, Pentalophos, Tsotyli and Ondria, lack equivalent structures indicating that they did not experience this deformation. The Middle to Upper Eocene period, related to the deposition of Krania Fm., corresponds to the first tectonic event (T1). It produced folds (verging towards NE) and reverse faults (striking WNW-ESE) which form the main compressional structures within the Krania Formation (Fig. 41). T1 event resulted in the deformation and uplift (emersion) of the Eocene strata during Upper Eocene. The uplift at the Krania sub-basins is deduced by the fact that the Eocene formation is covered in places unconformably by Uppermost Eocene, Oligocene and Miocene deposits (Fig. 26). Strike-slip faults, mostly with reverse component (i.e., in general NW-SE trending dextral faults and ENE-WSW sinistral faults; Fig. 41), which were described earlier (Chapter ) along the margins of Krania sub-basins, are assigned to this event, as they exhibit the same kinematic geometry and are related to the development of Krania sub-basins. Palaeostress analysis for T1 event indicates an almost horizontal maximum principal stress axis (ζ1), trending NE SW, and a sub-vertical minimum principal stress axis (ζ3), trending WNW-ESE (Fig. 41; Appendinx B). 102

131 Fig. 41. Northern Krania sub-basin, where the strike and dip of Krania strata is shown, as well as the observed structures at the margins and within Krania Fm., related to T1 event. The lighter yellow colour represents the continuation of Eocene Krania Fm; it is unknown which exactly is its extend underneath the younger deposits. Palaeostress analysis diagrams from different locations show fault planes, slip lines and principal stress axes (lower hemisphere, equal-area stereographic projection): big red circle, ζ1; medium orange circle, ζ2; small yellow circle, ζ3; single-direction arrows were used for normal and reverse faults, while double-direction arrows were used for strike-slip faults. Numbers on diagrams refer to Appendix B. The often occurrence of strike-slip movements with reverse component, especially along the western margins of the Krania sub-basins, related to positive flower structures and controlling the subsidence and organisation of the sub-basins, as well as the geometry of the compressional structures within the Krania Fm. strata, point to 103

132 a transpressional regime during Eocene, rather than a clear compressional one. The compressional structures (reverse faults and folds) are orientated sub-parallel to the main marginal dextral strike-slip faults orientation, which is an indicative geometry between a strike-slip zone and reverse structures for basins developed under transpressional regime (e.g., Busby and Ingersoll 1995, Einsele 2000). T2 event. The deposition during Oligocene was controlled by strike-slip faulting, as indicated by the occurrence of dextral strike-slip of NW-SE to NNW-SSE orientation between the basement rocks (Cretaceous limestones and ophiolites) and the Lower Oligocene strata of Eptachori Formation. Such faults also affect the Eptachori and Pentalophos Fms, as well as the older Krania Fm., on contrary to the younger Tsotyli Fm. which doesn t appear to be affected, and therefore assigned to a second tectonic event (T2), dating from Lower to Upper Oligocene. Dextral strike-slip faults of general NW-SE orientation occur along the western MHT boundary forming a steep morphology and also in Theotokos and Vassiliki villages area (centre and southern part of MHT). The strike-slip faults exhibit both normal and reverse dip-slip components (Fig. 42). At the Spileo and Theotokos villages areas, they trend parallel to each other forming positive flower structures that controlled the deposition during Oligocene time. The same apparently happens at Vassiliki village area, since the same structural pattern with this in Theotokos area occurs at the exact continuation of Theotokos area along NNW-SSE orientation (Fig. 42). Sinistral strike-slip faults striking NE SW to ENE WSW occur in many parts of the MHT with the same kinematic relations and relative ages as the NW SE dextral strike-slip faults, and therefore are interpreted as antithetic Riedel faults to the main dextral faults (Fig. 42). An equivalent example is the sinistral fault at Trikomo village, between Krania and Eptachori Fms. The reverse faults trending from NW-SE to W-E, which occur mainly in Eptachori and lower Pentalophos Fm., exhibit the same kinematic geometry with the strike-slip faults of T2 event and thus are assigned to the same event. Palaeostress analysis indicates an almost horizontal maximum principal stress axis (ζ1), with a NNE SSW strike (Figs 42, 43a, b), showing a small change in orientation compared with the first tectonic episode (T1; see Fig. 41, Appendix B). In contrast to ζ1, which retains almost the same orientation with T1 event, the minimum principal stress axis (ζ3) appears horizontal with an ESE WNW orientation (Fig. 43a,b). 104

133 Fig 42. Palaeostress analysis diagrams for structures of Eptachori and Pentalophos Fms related to T2 event (black) and T3 event (blue). Diagrams show fault planes, slip lines and principal stress axes (lower hemisphere, equalarea stereographic projection): big red circle, ζ1; medium orange circle, ζ2; small yellow circle, ζ3. Normal and reverse faults are shown with singledirection arrows; strike-slip faults are shown with double-direction arrows, black or white depending on the relative normal or reverse component respectively. Numbers on diagrams refer to Appendix B. The Eptachori and Pentalophos formations are shown on the map, where the dashed-lined area indicates the possible extension of the basin and depositional area during Oligocene - Lower Miocene. 1, 2, 3: strike-slip faults related to T1, T2, T3 events respectively; 4: thrusts/ reverse faults. 105

134 Evidence, as the small reverse component in the strike-slip faults, the occurrence of positive flower structures related to the basin s development, and the contemporaneous development of few compressional structures (reverse faults and gentle folding, more often in Eptachori Fm.), moreover with a sub-parallel disposition to the main dextral strike-slip faults, implies the strike-slip faults of T2 event developed under a transpressional regime during Oligocene times. The occurrence of strike-slip faults in Krania Formation related to T1 event and in Eptachori and Pentalophos Fms related to T2 event, as well as the palaeostress analysis and the inferred traspression for both periods, shows that the main tectonic regime continued from Eocene to Oligocene, changing in intension and only a little bit in orientation of the ζ1 stress axis. T1 event is characterized by more intensive compressional tectonics than T2 event, as evident by the higher frequency of compressional structures (reverse faults and folds) and the uplift of the Eocene subbasins which appoints the end of the first episode with a paroxysmic phase (it is marked by the unconformably overlying Uppermost Eocene strata of Eptachori Fm. on the Krania Fm. deposits). Krania sub-basins have been affected by the waning Eocene orogenic movements. In contrast, compression fades during the T2 event from Lower to Upper Oligocene, affecting stronger the Eptachori than the Pentalophos Formation. T3 event: The sinistral faults of NW-SE orientation, observed in few places along the western MHT boundary (i.e. Eptachori and Mesolouri (north from Filippei ) villages areas), show different kinematics than the dextral faults developing along the same orientation which dominate in the broader area of the western boundary (Fig. 42). This fact designates the sinistral faults of NW-SE orientation rather as particular, restricted in area, segments, in comparison to the major strike-slip fault along the western boundary of MHT. The sinistral strike-slip faults of this orientation are related to a WNW-ESE principal maximum stress axis (ζ1), in contrast to the dextral faults which are related to a NNE-SSW ζ1 stress axis (Fig. 42). However, the broader area lacks important structures resulting by a NW-SE orientated ζ1 stress axis, and such structures are also not documented for the Pindos thrust zone area. The only other comparable kinematically structures (Fig. 42, blue diagrams) are: (i) the lateral reverse fault with sinistral component of WNW-ESE 106

135 orientation cutting Eptachori Fm., at the Eptachori village area (number 11 in Fig. 29), (ii) the syn-sedimentary reverse faults cutting Pentalophos Fm. at the center of the trough (number 12 in Fig. 29), and (iii) the few NE-SW trending reverse faults indicated on the geological map sheet Grevena of I.G.M.E. (1972) to cut the Eptachori and Pentalophos Fms, which appear the same kinematic geometry. Despite the small number of existing structures registering the impact of WNW- ESE compressional tectonics, still this cannot be neglected or not recognized as a separate T3 event. This is certainly prior to the Tsotyli Formation deposition, because it is documented in Eptachori and Pentalophos Fms, but not in the any younger deposits. It seems to be a rather short event in time and of local importance, since it is detected only in very few locations, in contrast to T2 event. Its imprint on the Upper Oligocene Lower Miocene strata of Pentalophos Formation (i.e. even showing synsedimentary activity), and the striations observed in Mesolouri area, showing the dextral movement along NW-SE orientation preceding the sinistral one, ascribe this event at the end of Oligocene-beginning of Miocene (Fig. 42). Moreover, the occurrence of compressional structures related to a NNE-SSW orientated maximum principal stress axis (i.e. T2 event) in Eptachori and lower Pentalophos Formations (i.e. Oligocene times), is in accordance with the change of the orientation of the maximum ζ1 stress axis from NNE-SSW to WNW-ESE at the end of Oligocene. Palaeostress analysis shows that T3 event (Fig. 43c, Appendix B) is characterized by a horizontal/ sub-horizontal maximum principal stress axis (ζ1) trending WNW- ESE, and a sub-vertical minimum principal stress axis (ζ3) trending NNE-SSW. The intermediate principal stress axis (ζ2) appears sub-horizontal with a NNE-SSW orientation. Fig. 43. Palaeostress analysis diagrams for T2 (a, b) and T3 (c) events (see Fig. 42 for explanation of symbols). 107

136 T4 event. The forth tectonic event (T4) is related to the low-angle normal faulting with small strike-slip component along the eastern MHT boundary. These faults are assigned to the Lower Miocene, as they cause to the subsidence of the eastern part of the basin where the Tsotyli Formation was deposited. Paleostress analysis shows that ζ3 was almost horizontal with a NE SW orientation, ζ2 was sub-horizontal with NW-SE orientation, and ζ1 was sub- vertical; they indicate a generally extensional stress regime (Fig. 44, Appendix B). Fig. 44. Palaeostress analysis diagrams for low-angle normal faults observed along the eastern MHT boundary (bold black line), between the basement rocks (i.e., Pelagonian microntinent and superimposed ophiolites) and the Miocene Tsotyli Fm. (T4 event). Diagrams illustrate fault planes, slip directions and the main stress axes: (lower hemisphere, equal-area stereographic projection): big red circle, ζ1; middle orange circle, ζ2; small yellow circle, ζ3. Numbers on diagrams refer to Appendix B. 108

137 T5 event. The dextral strike-slip faults trending NNE-SSW cutting the Tsotyli and Ondria formations (near villages Nestorio, Tsotyli and Taxiarchis), but not the younger deposits, are inferred to relate to a compressional event (T5) of relatively local importance. The NW SE-trending reverse faults that cut Miocene strata of the Tsotyli Formation exhibit the same kinematic geometry and age, and therefore are regarded to have developed under the same stress regime, contemporaneously with the NNE-SSE trending dextral strike-slip occurring in the Lower-Middle Miocene formations (Fig. 45, 46). In addition, in the cross-section of the IGME Knidi Sheet (Mavridis et al. 1985), reverse faults of NW-SE orientation, are shown between the basement and the Tsotyli Formation east of Grevena. The T5 event is assigned to an Upper Miocene age, as it affects the Tsotyli and Ondria Formations but not the Pliocene and younger deposits. Paleostress analysis indicates that those faults were developed by an almost horizontal NE-SW trending ζ1 and an almost horizontal NW-SE trending ζ3 (Fig. 45, Appendix B). Fig. 45. Palaeostress A comparable compressional event of the same period analysis diagram for was documented by Kilias et al. (2001) further north, reverse and strike-slip within Albania, where the Mirdita ophiolites overthrust faults (T5 event; see Fig. Upper Miocene molasse-type sediments along reverse 40 for explanation of faults, giving rise to similar kinematic features to the T5 symbols). event. T6 event. The last tectonic event (T6), assigned to post-upper Miocene time, comprises high-angle normal faults that affect all of the MHT formations (M. Eocene - M. Miocene) and younger deposits. The high-angle NW SE to NNE-SSW trending normal faults within the MHT formations (south of Trikomo and Theotokos areas) and at the basement rocks around MHT margins, common between the Triassic carbonates of Pelagonian microcontinent and the post-miocene deposits, are assigned to this T6 event (Fig. 46). NE-SW to ENE WSW-trending normal faults of this stage are responsible for the elongate topography that is today followed by the Aliakmonas (north of the Pliocene Karperou Basin), Ionas (southern of Theotokos village) and Pineios Rivers (in Trikala 109

138 Fig. 46. Palaeostress analysis diagrams for the structures developed in the whole area of MHT after the Early-Middle Miocene times, related to T5 and T6 tectonic events; lower hemisphere, equal-area stereographic projection. Fault planes, slip lines and main stress axes are shown: ζ1= big red circle; ζ2= medium orange circle; ζ3= small yellow circle; single-direction arrows= normal and reverse faults, doubledirection arrows= strike-slip faults, black or white depending on the relative normal or reverse component respectively. Diagrams in blue colour indicate the fault sets related to the compressional T5 event. The orientation of the ζ1 and ζ2, or ζ3 and ζ2 axes are illustrated in each diagram with arrows at the perimeter of the diagram, with red and grey, or white and grey colour respectively. Numbers on diagrams refer to Appendix B. 110

139 region, not comprised in Fig. 46). The orientations of these faults, as well as this of faults trending NW-SE, suggests that at least some of the normal faults utilized preexisting weak zones of Oligocene age at a time when strike-slip faults of the same trend developed under a different stress regime (i.e. T1 and T2 events). It is also possible that some rivers were controlled by even earlier structures. The Plio- Pleistocene age and the rectangular shape of the small Karperou basin (directly south of the Vourinos Massif and the Aliakmonas River) reflect a possible control by young ENE WSW normal faults (at its northern and southern margins); these faults were active after the end of the Miocene, causing subsidence of specific areas. It is noteworthy that the Karperou basin s northern margin coincides with the Aliakmonas River (Fig. 46). Palaeostress analysis for the majority of high-angle faults of T6 event indicates the minimum principal stress axis (ζ3) oriented horizontally from NE SW to ca. N-S; ζ1 is vertical, while ζ2 axis is usually also horizontal striking NW-SE (Figs 46; 47b, c; Appendix B). However, few cases of high angle normal faults denoting an almost horizontal NW-SE trending ζ3, a sub-vertical NNE-SSW trending ζ1, and a subhorizontal NE-SW trending ζ2, also exist (Figs 46, 47a, Appendix B). The later palaeostress analysis concerns normal faults that were all observed in Miocene but not any younger deposits; this is an evidence of dating the development of those faults in Upper Miocene-beginning of Pliocene, exactly after the compressional T5 event. Still, a more detailed study is needed in order to support with confidence this assumption and we refrain from distinguishing it as a separate event. Some of high-angle post Miocene normal faults show a recent activity; they relate to a ζ3 of N-S to NNW-SSE orientation and a vertical ζ1, which are compatible with the present active tectonics of the area, computed by the focal mechanisms (Papazachos & Papazachou 1997). The recent seismic activity also produced an earthquake in the area in 1995, of magnitude 6.5 Richter (Mountrakis et al. 1998; Pavlides and Mountrakis 1987; Tranos 1998, 2009). This earthquake caused great damage to the city of Kozani and many villages in the area. The T6 event marks a generally extensional period, which initially started in the Lower Miocene and continues today (i.e., ζ3 is north south). 111

140 Fig. 47. Palaeostress analysis diagrams for high-angle normal faults (T6 event; see Fig. 46 for explanation of symbols). a. ζ3 trending NE-SW (related to the extensional regime dominated after Miocene times); b. ζ3 trending NW-SE (normal faults possibly developed during Upper Miocene-beginning of Pliocene); c. ζ3 trending N-S to NNW-SSE (active tectonics). 112

141 7. SUBSIDENCE RATES OF THE MHT It is well known that the initial subsidence of a basin occurs due to isostatic compensation because of changes in thickness of the layers in the lithosphere. The elevation of the top of the crust is a function of the thicknesses and densities of several layers. Assuming a horizontal datum surface in the viscous mantle asthenosphere, the mass of the overlying rock column per unit area (i.e. solid upper mantle lithosphere, solid crust consisted of igneous and metamorphic rocks, sediments and water) remains constant. Hence, local or regional load of the lithosphere leads to the flexural response of the crust and subsidence; similarly, unloading leads to some regional uplift. Then, as a basin is covered by water and starts filling with sediments, the sediment and water load cause additional subsidence. In order to determine the tectonically driven part of subsidence of a chosen location in a sedimentary basin, the technique of back-stripping is used. Main purpose of this technique is to calculate and remove the effects of sediment loading, changing palaeo-bathymetry, the sea-level variations and the compaction of the sediments (Watts and Steckler 1979, Grandstein et al. 1985, Steckler et al. 1988, Einsele 2000). First step of back-stripping is the decompaction of different units of a sediment section; it has the meaning of reconstructing the original sediment thickness. For its calculation, the present thickness and mean porosity of a sedimentary column are needed, as well as the original mean porosity which can be obtained by the appropriate porosity-depth curves, determined for particular lithologic units. Next step of this technique is the isostatic adjustment, which aims to find out the tectonic or thermo-tectonic subsidence of the basin which would have taken place without the sediment load. For this, the loads of the various sediment units are removed. Further calculations of back-stripping comprise the removal of the effects of the paleobathymetry and sea-level changes. Different processes controlled by eustatic sealevel rise or fall, in combination with the increased or reduced sediment accumulation, or even sediment erosion, can lead to different results. Subsidence can be calculated by the local Airy-type isostatic adjustment with the appropriate corrections related to 113

142 the particular conditions of eustatic sea-level changes and sedimentation rates accumulation, or even intermittent sediment erosion. As resulting from the above brief description of the back-stripping method, the determination of subsidence history of a sedimentary basin in certain locations composes a special study, which includes several steps and requires the consideration of many different parameters. In this technique, the sedimentary section of each location and the underlying basement should be well known. It comprises the knowledge of formations and basement thicknesses, the type of sediments, the bio- and chrono-stratigraphy, the paleo-environment and the paleowater depths, and the relative sea-level changes. The great depth that the sedimentary formations of the Mesohellenic trough reach, as shown by the iso-depth contours in figure 48, and the large volume of sediments accumulated from Eocene till Miocene, show that the MHT experienced periods of significant subsidence and high sedimentation rates. In this study, an approximation of the basin subsidence and uplift through time in 12 different locations of the Mesohellenic trough has been attempted, using the present depths where basement is detected via geophysical data (Fig. 48), the known thicknesses of the sedimentary formations, the documented water-depths for the different formations, the eustatic sea-level changes (i.e. eustatic curve from Haq et al. 1987), and the present altitude of each location. Decompaction of sediment units and isostatic compensation adjustments were not made, as the exact stratigraphy and the basement thickness were not known; hence, errors are induced in the overall subsidence curves. Lack of decompaction equals to lesser amounts of subsidence comparing to the original one. Overall, the available information was not adequate for an accurate and more detailed study, but none the less, an idea of the approximate subsidence history was succeeded. The curves for the subsidence of the basement (Fig. 49 A to C) were drawn for the time-period when sedimentation was taking place at each specific location, while a possible curve up to the known present-day situation is given within a tinted area. The continuation of the curve (tinted area) is based on the subsequent to sedimentation sea-level changes at each location, regarding there was no more sediment accumulation, but in contrast some erosion providing the active depocenters with reworked sediments (~0 to 200m decrease of sediments thickness 114

143 was arbitrary attributed to erosion at each location after the end of sedimentation). The coloured area represents the errors related to the time and rate of uplift of the each location over the sea-level surface, up to present altitude. Fig. 48. Iso-depths (depths in meters; v= 4km/sec; modified after Zelilidis et al.2002) on top of the MHT geological map. The locations for which the subsidence curves are given in figure 46 (A to C), are shown. 115

144 Fig. 49. A. Area of Krania sub-basin and east of it. Subsidence-uplift rate curves for certain locations shown in figure 48. Red curve of each diagram indicate the basement subsidence, taking under consideration the sea-level changes, and thus is closer to the true amount of subsidence than the black curve, from which the sea-level changes are excluded. The tinted area corresponds to the possible depth that basement ranged at each location after the end of sedimentation until present day situation (see text for explanation). 116

145 Fig. 49. B. Area along the western MHT margin. 117

146 Fig. 49. C. Area along the eastern MHT margin. 118

147 The curves of figure 49 show rapid subsidence and uplift of different parts of the Mesohellenic trough during different periods of the MHT evolution. The calculated subsidence and uplift rates correspond only to the average of mean values of different locations of the MHT area, during the depositional period of each formation. They do not provide information about the higher or lower rates that probably took place during each period; this would demand a detailed bio- and chronostratigraphical record which was not available. Nevertheless, a basic idea of the subsidence framework is given. The localized higher or lower mean subsidence rates are illustrated in the diagrams of figure 49. Furthermore, small subsidence and uplift indicated by the curves may correspond to sea-level fluctuations or erosion of deposits, without any vertical movements of the crust taking place. The primary depression developed at the Krania sub-basin area during the Middle- Upper Eocene, exhibiting a subsidence of ~ m/ Ma, before the uplift and emersion of this part, which should have been the minimum 400 meters. Afterwards, during the end of Eocene-Lower Oligocene related to Eptachori Fm. deposition, the curves manifest that the subsidence was enhanced along the western part of MHT, with an average rate of ~275m/Ma and reaching a maximum of ~335m/Ma south of Pentalophos village (see Figs 48, 49B; number 7). This period was followed by rather smaller subsidence rates during Upper Oligocene-beginning of Miocene when Pentalophos Fm. was deposited, characterized by an average of ~220m/Ma and reaching highly different maximum values between ~310 and ~360m/Ma at the places where the biggest depths of the MHT were found (Figs 48, 49B, C; numbers 7, 8 and 10). Finally, the last Lower-Middle Miocene period of Tsotyli and Ondria Fms deposition, was characterized by the lowering of the easternmost part of the trough, with similar to the previous period subsidence rates of the same average (i.e. ~220m/Ma), while reaching higher rates of ~300m/Ma in places (Figs 48, 49C; numbers 11 and 12). Regarding the present altitude of the MHT area, which ranges between ~500 m in the south to ~1300m in the north, it is easily inferred that the whole area was elevated after Middle-Upper Miocene (i.e. end of MHT deposition) at least ~400 to more than 1200m, including the sea level fluctuations which correspond to an overall drop of about 100m since the Middle-Upper Miocene time. This uplift is easily deduced by the diagrams showing the subsidence uplift rates at certain locations of the MHT during its evolution. 119

148 120

149 8. STUCTURAL EVOLUTION DISCUSSION The formations of MHT were deposited sub-parallel to one another, showing a migration of depocenters from west to east through time. The major geodynamic control of the basin s development is attributed to underthrusting of the External Hellenides beneath the Pelagonian microcontinent (e.g., Ferrière et al. 2004, Vamvaka et al. 2006). The thrust load of the Pindos zone, together with the sedimentary load of the accumulated clastics on top of the External Hellenides during Oligocene (i.e. Ionian-Gavrovo basin at the west of MHT and Pindos trust zone; see Fig. 1 for location) is considered to have played an important role in the enhanced subsidence and underthrusting of External Hellenides (Sotiropoulos et al. 2003). We have recognized five main stages in the evolution of the MHT region in relation to regional tectonic events (Figs 50 to 57). The basin evolution was initiated in the Middle Eocene, nearly contemporaneously with the thrust-imbrication of Pindos units (Jones & Robertson 1991) and the deformation of the Pelagonian upper plate (Mountrakis 1986; Kilias et al. 1991a). A transpressional regime with the maximum principal ζ 1 stress axis orientated NE-SW dominated this period (T1 event; Fig. 41). The first sub-basins comprising Eocene deposits developed on ophiolitic basement by flexural subsidence beneath the advancing load of the thick Pindos thrust-fold belt (Fig. 52; for examples of flexural processes see e.g.: Moxon and Graham 1987, Karner and Watts 1983, Royden and Karner 1984; Einsele 2000). Passive isostatic subsidence became active as the basin began to infill with sediments. Complicated stratal relationships developed in places between the basin fill and the evolving structural high above the Pindos accretionary complex. Dextral strike-slip faults of ophiolitic rock assemblage at the western sub-basins margins have also affected the subsidence of the Eocene subbasin at the west. They are described by a reverse component and are related to a transpressional regime (Figs 53, 54). A shear character is evident at the contact zone of the basement rocks with Krania Fm. strata which appears almost vertical and with a characteristic waving-shape along the bedding. In contrast to the intensive deformation apparent at the western Krania sub-basins margins, compression was not that strong at the eastern Eocene sub-basins margins. Transgressive basal conglomerates are evident at the eastern margin of the southern Krania sub-basin, whereas there is no tectonic contact (Fig. 53). 121

150 Fig. 50. Geological map of the MHT with all faults relating to the evolution of the trough from Middle Eocene onwards. 122

151 Fig dimensional portray of the geological map of MHT. Morphology, in many cases, follows and emphases the tectonic lines and structures (such as the folding and curve of the folding in western Krania sub-basin; the high inclination of Eptachori strata lying on the faulted western MHT boundary basement rocks; the uplifted flower-shape structure in the southern Theotokos-Vassiliki area; the high-anlge normal faults along the eastern MHT boundary at the north). The legend of the geological units is given in figure

152 Fig. 52. Schematic cross sections and map view showing the first stage of the MHT tectonic evolution during Middle Upper Eocene (T1 event) (Vamvaka et al. 2006). 124

153 At the northern Krania sub-basin, the strike of the strata and the orientations of the b-axis of the folds (i.e. NW-SE orientation in the centre and curving towards W-E at the northern-edge, and towards SW-NE at the southern- edge of the northern Krania Fm. outcrop) coincide with the orientation of the sub-basin margins, and the shape of the sub-basin (Fig. 26). The latter implies an heterogeneous promotion of the different parts of the ophiolites and the deposited sediments towards east; eastward verging of the folds is in accordance with this movement. Transverse faults with ENE- WSW trend and lateral slip along the northern Krania sub-basin margins, may also relate to this movement, and thus have occurred during the same period, playing role to the basin s organisation. The first stage ended in the Upper Eocene with the uplift and deformation of the two Eocene (Krania) sub-basins (Figs 52 to 54). This deformational stage precedes the deposition of Uppermost Eocene alluvial fans, since they rest unconformably on the almost vertical Krania strata and without showing any similar structure, as well observed south of Trikomo village (Fig. 26). The FT data record the main Eocene orogenic event in western Pelagonian microcontinent a little earlier, during Early-Middle Eocene times, directly followed by high exhumation rates. This is explained by the migration of the main compression to the west through time, affecting strongly and uplifting the Krania sub-basins in the Upper Eocene. The second stage is associated with the subsidence of an elongate, relatively narrow basin in which the Eptachori and Pentalophos Formations were successively deposited from Lower Oligocene to Lower Miocene time (Figs 42, 50). This stage was controlled by strike-slip faults bounding the basin, while the main contraction during continuing convergence migrated westwards (Figs 53 to 56). In this sense, the basin should be seen as a type of strike-slip basin between strike-slip faults. The strike-slip fault at the western margin of the basin acted as the master fault controlling the basin development, whereas the strike-slip fault that probably bounded the eastern margin of the basin, was covered by the Tsotyli Formation (Fig. 55, 56). The strike-slip faults developed at the centre and the southern part of the trough (i.e. Theotokos and Vassiliki village areas; Fig. 50) constitute a long continuous strike-slip zone that apparently induced subsidence mainly along its western side during Oligocene, whereas it probably bounded in a way the basin to the east at this region (Fig. 53A, B). 125

154 Fig. 53. Cross-section across the MHT, from mountain Koziakas towards the Pelagonian microcontinent to the east, passing north of Vassiliki village (cross section AA in figure 50). 1. Pelagonian microconent gneisses and schists, and the overlying Triassic limestones; 2. Middle-Upper Jurassic limestones (Koziakas sequence), deposited at the eastern part of the old Pindos basin (referred as Hyper-Pindic zone in geological map sheet Kalambaka of I.G.M.E., 1972); 3. Tythean ophiolites; 4. Cretaceous limestones, overlying the ophiolitic assemblage; 5. Krania Formation (Middle-Upper Eocene); 6. Eptachori Formation (Lower Oligocene); 7. Pentalophos Formation (Upper Oligocene-Lower Miocene); 8. Tsotyli Formation (Lower to Middle Miocene); 9. Holocene deposits; 10. Dextral strike-slip faults of general NW-SE orientation (Upper Eocene-Oligocene times); 11. Normal faults (low- and high- angle is showed by the inclination of the fault lines, corresponding to Lower Miocene and post-miocene times respectively). Open to darker tones of grey-black colour are used for the faults according to the time of their development from older to more recent times (see text for explanation of tectonic events and evolution). 126

155 Fig. 54. Cross-section across the MHT, passing from Grevena city area, where the trough has a depth of more than 4000 meters (see Fig. 4). (cross section CC in figure 50). 1. Pelagonian microconent; 2. ophiolites; 3. Cretaceous limestones, overlying the ophiolites; 4. Krania Formation (Middle-Upper Eocene); 5. Eptachori Formation (Upper Eocene-Lower Oligocene); 6. Pentalophos Formation (Upper Oligocene-Lower Miocene); 7. Tsotyli Formation (Lower to Middle Miocene); 8. Pliocene-Quaternary deposits; 9. Dextral strike-slip faults of general NW-SE orientation (Eocene-Oligocene times); 10. Normal faults (low- and high- angle is showed by the inclination of the fault lines, corresponding to Lower Miocene and post-miocene times respectively). Open to darker tones of grey-black colour are used for the faults according to the time of their development from older to more recent times (see text for explanation of tectonic events and evolution); reverse faults developed during Middle-Upper Eocene times and affecting the Krania Formation are shown too, with white colour. 127

156 Fig. 55: a, b. Caption and legend in the next page. 128

157 Fig. 55. Cross-sections across the MHT, passing from Theotokos uplifted structure, and showing the evolution from Lower Oligocene to present time (from a to d; cross section BB in figure 50). a, b: the marginal strike-slip faults and the flower structure at Theotokos village area of T2 event; c. the low-angle faults of T3 event, causing the subsidence of the eastern part of MHT; d. the present day situation showing the last high-angle faults of T5 event; 1. External Hellenides; 2. Pelagonian microconent; 3. ophiolites; 4. Cretaceous limestones, overlying the ophiolites; 5. Eptachori Fm. (Lower Oligocene); 6. Pentalophos Fm. (Upper Oligocene-Lower Miocene); 7. Tsotyli Fm. (Lower-Middle Miocene); 8. Holocene deposits; 9. Dextral strike-slip faults of general NW-SE orientation (Oligocene time); 10. Normal faults (low- and high- angle is showed by the inclination of the fault lines, corresponding to Lower Miocene and post-miocene times respectively). Black colour is used for active faults, and more open colours are used to indicate successively older inactive faults. 129

158 Variable depths along the axis of the basin, as shown on a palaeo-bathymetry map (Fig. 48), and the uplifted structure in the middle of the basin, can both be related to strike-slip faulting. Strike-slip basins commonly experience localized episodes of rapid subsidence or uplift, resulting in unconformities (Fig. 49, 54, 55a,b, Pict. 44). The synclinical structure formed by the Eptachori and Pentalophos Formations (Figs 50, 55) can be related to syn-sedimentary tectonics and the continuing sedimentary loading of the area during the second stage (T2). Pict. 44. Unconformity in the Pentalophos formation strata at the northern part of MHT (west of Nestorio village), showing sedimentary and probably tectonic instability during the deposition (i.e. Upper Oligocene times Lower Miocene times). Alternative models for the evolution of the MHT, relate the main subsidence of the trough to reverse (Doutsos et al. 1994) or normal faulting (Ferrière et al. 2004). Doutsos et al. (1994) relate the evolution of the MHT to compression and backthrusting. However, no clear compressional structures have been identified in the field, specifically thrusts dipping to the west (except for the Krania Formation). By contrast, Ferrière et al. (2004) relate the main subsidence of the basin to flexures caused by underthrusting and normal faulting after Lower Oligocene. Although, normal faulting is recognized here too, this is inferred to after the beginning of Miocene. 130

159 Fig. 56. Schematic cross section and map view showing the second stage of the MHT tectonic evolution during Oligocene Lower Miocene (T2 event) (Vamvaka et al. 2006). 131

160 The maximum principal stress axis (ζ 1 ) shows only a small change from its previous orientation during Eocene (i.e. from NE-SW towards NNE-SSW), whereas by contrast the extensional stress axis (ζ 3 ) was almost horizontal (T2 event; Fig. 43a, b). A transpressional regime is presumed by the small reverse dip-slip component of many strike-slip faults and the occurrence of positive flower structures (Figs 53 to 56; i.e. at Theotokos and Vassiliki villages areas at the central and southern parts of the MHT, and at Spileo village area at the western MHT margin); few reverse faults also occur, sub-parallel to the main dextral strike-slip faults. As described previously, the stress regime didn t alter a lot since Eocene times. In contrast, after the paroxysmic event in the Upper Eocene that resulted to the uplift and deformation of Krania subbasins, almost the same stress regime continued to affect the area, but with less intension, as deduced by the rareness of compressional structures (reverse faults and folds) in Eptachori and Pentalophos Formations comparing to those within the Krania Fm.. At the end of second stage of MHT evolution, during the end of Oligocene beginning of Miocene, a short-time deformation took place (T3 event), resulted by compressional tectonics and related to a maximum principal stress axis along WNW- ESE orientation. Sinistral strike-slip faults trending NW-SE in places at the western boundary are attributed to this deformational event (T3), as well as few reverse faults of NE-SW strike in one location in Pentalophos Formation (Fig. 42, 43c). The scarceness of tectonic structures related to this stress regime implies it had a local influence. A short time event with different dynamics from the governing one of the second stage of MHT evolution (i.e. T2 event) could be adequate to cause a change in the sense of movement along pre-existing weak fracture zones. Therefore sinistral strike-slip faults were developed, at least on specific segments, along the pre-existing dextral strike-slip fault zone along the western boundary of MHT. Older conjugate reverse faults that cut such sinistral faults near Eptachori village, at the ophiolites and syn-tectonised Upper Jurassic to Lower Cretaceous limestones near the contact with the Eptachori Formation, must have developed between Upper Cretaceous and Lower Eocene times, since they do not seem to relate with the evolution of MHT and taking under consideration the age of the limestones. The next deformational stage (T4 event) involved low-angle normal faulting, and it took place from Lower to Middle Miocene (Figs 50, 55c, 57). Sub-horizontal NE-SW orientated extension dominated this period (Fig. 44). The compression moved further 132

161 west and the region experienced late orogenic collapse under a plate convergence regime (Kilias et al 1991a,b; Mountrakis et al. 1992) (Fig. 57). This change in tectonic framework resulted in deformation of the eastern margin of the inferred Oligocene strike-slip basin. The marginal low-angle NW-SE normal faults with a small lateral component cut the previous eastern margin of the MHT and caused subsidence and some widening to the basin where the Tsotyli and Ondria Formations were deposited (Figs 50, 53-54, 57). Low-angle normal faulting used pre-existing Oligocene/Miocene low-angle extensional shear zones of similar kinematics, which also affected the surrounding ophiolitic and Pelagonian basement rocks (Kilias et al. 1991a,b) (Pict. 34 and Fig. 42). At the southeastern-most part of the MHT, the Tsotyli Formation was deposited directly on Eocene strata (Fig. 50, 53), indicating that this specific region was not an active depocenter during Oligocene Lower Miocene; this was apparently caused by the uplift and deformation of Krania sub-basin at the end of Eocene. This area started receiving again sediments since Lower Miocene times, when low-angle normal faulting affected the area. The deposition of the Ondria Formation followed during Middle Miocene times in rather shallow water. Today the latter formation remains only in a few places of the MHT (Fig. 50). AFT data from the sediment samples of the MHT deposits register the extensional period of Lower Miocene with the abruptly decreasing lag-times of the young AFT age groups of the samples with Lower Miocene stratigraphical age [i.e. lower stratigraphical levels of Tsotyli Formation (sample AV-101) of Aquitanian/ Burdigalian age (~20Ma); Fig. 19]. Moreover, the broad track-length distribution and the modelling of the basement sample AV-23, with AFT age of ~25Ma, intimate a reheating or at least a prolonged stay in the APAZ around Lower Miocene that can be related with the extension and crustal thinning of the area during this period. The increasing lag-time during Middle Miocene that followed (Fig. 19, sample AV-107) is in consistence with the final filling of the MHT and the finishing of its evolution. During the Upper Miocene, the compressional T5 event caused strike-slip faults and local thrusting of the MHT and in some places overthrusting of the ophiolites to Miocene sediments. T5 compression occurred in the midst of a generally extensional period, characterized by orogenic collapse and uplift of the Hellenides after Eocene crustal overthickening (Lister et al. 1984; Kilias et al. 1991a,b, 2001). This shows that orogenic extension can occasionally be interrupted by compressional events. 133

162 Figure 57. Schematic cross section and map view showing the ultimate stages of the MHT tectonic evolution from Lower Miocene to present time (T4 and T6 events) (Vamvaka et al. 2006). 134

163 Finally, the late stage of the MHT evolution is connected with the T6 deformational event which affected the MHT and the younger deposits from the Upper Miocene to present day. This period is characterized by high-angle normal faulting, showing variable orientations (Figs 46, 50). The normal faults of NNW-SSE orientation developed in Theotokos area, which developed in this stage, have firstly thought to account for the direct contact of Tsotyli with the Eptachori Formation, without the interposition of Pentalophos Formation (Vamvaka et al. 2006). However, a partial emersion of the specific region, where Eptachori Formation outcrops today had probably taken place, constraining the sedimentation during Lower Miocene at the western side of the uplifted domain; this was caused by the development of the positive flower structure due to strike-slip faulting of the second stage of the MHT evolution during Oligocene (Fig. 50, 55a,b). Therefore, the Pentalophos Fm. may have been eroded or never deposited at the specific region, and consequently the Tsotyli Fm. could directly overlie on the existing Eptachori Fm (Figs 55c). The high angle normal faults at this area cut all occurring MHT formations (i.e., Eptachori, Pentalophos and Tsotyli), and they can relate to reactivation of the former strike-slip faults of the second stage, described by the same NNW-SSE orientation (Figs 55d, 57). Generally, some of the older weak fracture zones of NW-SE and NE-SW orientations were re-activated during the extensional post-miocene period as normal faults. In the end, some of the normal faults related to the last extensional period and a minimum principal stress axis of ca. N-S orientation are still active. An approximate pattern of the subsidence occurred in the MHT during the stages of its evolution is provided by the calculated mean subsidence rates for different location within the trough (Figs 49, 48). A general picture of the amount of subsidence is drawn in comparison to the different stages of MHT evolution, despite the fact that the inferred average subsidence rates (calculated for the depositional period of each MHT formation) correspond to constant mean rates during the whole time of each formation s deposition, without including the probable variations in the interim. An estimation of the amount of uplift was also succeeded, for the periods: (i) at the end of Eocene related to the deformation of Krania Formation, and (ii) after Middle Upper Miocene, for which uplift is presumed by the present day altitudes in the area of MHT. 135

164 As previously documented, the subsidence of the first Eocene sub-basins initialised due to flexural response to the advancing load of the Pindos thrust-rockpiles; sedimentary load and water load caused further subsidence due to isostatic compensation, whereas tectonics played also an important role. The average subsidence rate during Middle-Upper Eocene, related to the development of the northern Krania sub-basin was about 240m/Ma, while the uplift occurred in Upper Eocene was at least about 400m, regarding that the Krania Formation was emerged over the sea-level as evident by the reddish colour of the first late Eocene breccias and conglomerates which rest unconformably on Krania Fm.; the latter intimates a terrestrial environment. Subsidence intensified during Lower Oligocene with an average value of ~275m/Ma, while higher mean values of ~335m/Ma are also present. This is associated mainly to the interplay of tectonic activity of strike-slip faults, and also the increasing load of the accumulating sediments. During Upper Oligocene beginning of Miocene, the average rates decreased to an average of ~220m/Ma, but still reaching significant higher mean values locally, with a maximum of ~360m/Ma. The generally smaller subsidence rates are in accordance with the tectonic regime characterized by decreasing intension and a change at the end of Oligocene, while the presence of local episodes of enhanced subsidence are related to the strike-slip fault activity. Finally during the third stage of MHT evolution of Lower-Middle Miocene, mean subsidence rates increase at the eastern part due to low-angle normal, especially rather close to the eastern boundary with a value of ~300m/Ma. As isostatic compensation triggered the beginning of MHT development, the same process is presumed to have led to the emersion of the area of the trough. As the area of the trough was filled with deposits, the resulting mass occupying the subsided space was smaller than the mass of the rocks initially occurring at the same region (i.e. igneous, metamorphic and sedimentary rocks), since the density of the sediments is lower than this of igneous and metamorphic rocks. Consequently, in order to re-establish the equilibrium of the lithospheric columns at the area of the trough, it is reasonable that uplift of the MHT region took place due to isostatic compensation. The uplift resulting for the present altitudes of the MHT area ranges from 400 to more than 1200m during the last ~15-10Ma. The amount of elevation of each locality was probably depended on the former amount of subsidence and loss of total mass from the balanced conditions. Uplift of the MHT area after Middle 136

165 Miocene was enhanced due to the cotemporaneous regression of the sea during Tortonian and removal of the extra water load. The restricted occurrence of Ondria Formation may relate to the rapid uplift which must have started around the late Middle Miocene times regarding the age of the last MHT formations. Thermal modelling for the Pelagonian rocks, resulted from track length distributions, indicates an increased cooling rate around 10 Ma, which is in accordance with the geology of the area, the contemporaneous filling and uplift of the MHT area in the end of Miocene. 137

166 138

167 9. BASIN TYPE CLASSIFICATION OF MHT As it is already mentioned, the Mesohellenic trough developed on top of the Tythean ophiolites, between the External and Internal Hellenides, during their continuing convergence and underthrusting of the External Hellenides beneath the Pelagonian microcontinent. Taking that under consideration, the MHT should be associated and compared to basin types which develop at convergent plate margins, whereas also the transform settings cannot be excluded as they may relate to different tectonic regimes. 13 basin types accrue to be correlated with the case of MHT, as relevant to those settings (Figs 5, 58 and 59; Table 1); from those types, five can be left out directly as they are associated with active subduction processes (i.e. trenches, and trench slope basins, forearc, intra-arc, and backarc basins; Fig. 6, Table 1), while in contrast the evolution of MHT took place during collision of continental blocks. However, Papanikolaou et al. (1988) characterized the MHT as behind the trench accretionary prism. Their reference to the corresponding volcanic arc of this period, tracking it down in North Aegean, induces the idea of a fore-arc basin. Also Ferrière et al. (2004) characterizes the first Krania sub-basins as deep forearc basins related to Pindos basin subduction. If Pindos basin was indeed an oceanic basin, and there was an active subduction taking place during Eocene times, then the MHT could be thought as a forearc basin during the first stage of its evolution (Eocene times), considering that some of the Krania sub-basins characteristics fit to those of forearc basins; these are the sedimentary facies changing from deep to more shallow environments (turbidites to fan deltas), and the ophiolitic fold-thrust zone bounding the west side of the basin and supplying it with sediments. Despite those similarities to forearc basins, and besides the issue of the existence of an active subduction, still the volcanic arc which supposes to bound the eastern side of the basin, and which is a typical criterion for the identification of those basin types (Fig. 58), is missing. The volcanic activity in North Aegean during Oligocene-Miocene is placed much too far to the east to correlate with the MHT development. The other types left to correlate with the MHT are: (a) those related to collision processes, which are foreland (retroarc, peripheral, intermontane), remnant and piggyback basins, and (b) the strike-slip basins. 139

168 Fig. 58. Profiles of three orogen types related to subduction: intra-oceanic, between continental and oceanic crust, and intercontinental collision (Scholl, von Huene et al. 1980). The positions of the trench, the fore-arc, foreland and retroarc basins are shown. The term foreland is inherited from former to plate-tectonics times and is used to characterize the basins which develop between an orogenic zone and a cratonic area (Figs 58 and 59). MHT is located in a foreland area, developed during continuous continental convergence, behind the Pindos thrust zone. Thrust and sedimentary load led to extended lithospheric flexure in the beginning of MHT evolution during Eocene, which is typical characteristic of the foreland basins; compressional tectonics in the thrust zone is the primary cause of subsidence in foreland basins, while the sedimentary processes play a secondary but still significant role. Foreland basins are further distinguished in three categories, depending on their exact geodynamical position. The term retroarc (Fig. 58) was proposed by Dickinson (1974) for the foreland basins developed behind compressional arcs, on 140

169 contrast to the peripheral foreland basins (Fig. 59a) that developed on subsiding plates during continental collisions. Foreland intermontane basins may form between basement elevations in the retroarc foreland basins regions (Dickinson & Snyder, 1978), when low-angle subduction takes place under compressional trench-arc systems (Fig. 59a and b). From the three types mentioned above, peripheral basins can be excluded due to their position on the subsiding plate, in contrast to the MHT which is developed on the upper plate of the colliding continental blocks. Fig. 59. Tectonic classification of collision- related basins (see Table 1; convergent settings), for (a) partial collision of continents with irregular shapes and boundaries which do not dit each other, leading to crustal over- or underthrusting, and (b) A-subduction zone (after Ampferer, Alpine-type); depression and flexuring of the continental crust occurs (Einsele, 2000). 141

170 Between retroarc and intermontane basins, the second term matches better to the MHT, as it is exactly related to low-angle subduction and underthrusting. Moreover, retroarc basins are related to clear compressional tectonics, while the intermontane basins may comprise rotational deformation and develop in different shapes and with different characteristics (e.g., narrow and fault-bounded, with a through drainage and association with strike-slip faults; wide, equally extended, surrounded by basement elevations; elongated with asymmetrical synclines and an uplifted margin on one side; Busby and Ingersoll, 1995). Another type of clear convergence related basins is remnant basins; they are associated with inherited ocean morphologic depressions, shrinking between the colliding continental margins, and finally subducting or deforming within suture belts (Fig. 59a). If the existence of Pindos Ocean is accepted and therefore the ophiolites upon which the Mesohellenic Trough is developed are related to this ocean, then one could contemplate the possibility of the trough being a remnant oceanic depression after the closure of the ocean. Despite the query about Pindos Ocean existence, the whole MHT did not suffer strong deformation as would be typically expected for a remnant basin; those basins are normally fated to destruction within suture zones, which is not the case of MHT. Finally, piggyback basins are determined by Ori and Friend (1984) as basins developed and filled during transported on moving thrust nappes (e.g., Apennines, Pyreneans). They are dynamic structures where sediments are concentrated, originating from the related thrust zones which may be peripheral, retroarc or transpressional settings. This term has been used excessively by the different researchers to characterize the MHT, but for different reasons (see Ch.2.3). Wilson s (1993) is referred only to the northern Krania sub-basin as piggyback basin, and it seems reasonable, since the ophiolitic thrust system was active throughout Eocene times, and even later in more western regions. About the characterization of the whole MHT as a piggyback basin, Ferrière et al. (1998, 2004) argument is accepted (Vamvaka et al. 2006), since the trough was developed on the ophiolitic nappes during the eastward underthrusting of the Gavrovo-Tripolitsa platform under the Pelagonian microcontinent. Except the basins related to convergent settings, strike-slip basins are also examined, since the development of MHT is strongly related to strike-slip fault activity. There is a big variety of sedimentary basins related to strike-slip faulting and may 142

171 occur at different plate boundaries (Fig. 60). In general, different cases of sedimentary basins related to strike-slip faults may comprise: mismatching along the margins, through-going and lateral asymmetry, episodic rapid subsidence, abrupt lateral facies changes and regional unconformities, and intense differences in stratigraphy, facies geometry and unconformities between particular basins of the same area. In mature stages of evolution, new strike-slip faults may develop obliquely across the basins; this may result to a straightening of the principal displacement zone (e.g., Doodley and McClay 1997). Fig. 60. Types of strike-slip fault patterns and resulting strike-slip basins. A, B, and C, braided faults; D, fault termination; E, en echelon faults (Reading 1980). There are three basic types of strike-slip basins, depending on the subsidence mechanisms, which are briefly reviewed here; subsidence results either from crustal attenuation during and after extension, or from lithospheric flexure due to sedimentary load and compression. Therefore, transtensional basins form due to extension along the releasing bends of continuous through-going faults, and at the intersections of bifurcating faults or discontinuous fault segments (fig. 61A; Crowell 143

172 1974a, Mann et al.1983, Christie-Blick and Biddle 1985). Transpressional basins are related to compression and develop along steep restraining bends at intensively deformed and thrusted margins due to tectonic and sedimentary load (Crowell 1982; fig. 61B, 62A). Alternatively, transpressional basins may develop at the wedges formed by strike-slip faults at weak restraining bends; one or both faulted margins is/are uplifted, resulting to the subsidence of the middle area, as the one block moves and leaves behind the restraining bend (Crowell 1974a, b, 1982; Fig. 61B). The compressional component of the transpressional systems can be inferred from wrench faults and fold belts of limited extend (Fig. 12). Finally, the term transrotational is used for the basins that develop in strike-slip systems due to rotation of crust around vertical axes (Ingersoll 1988). Fig. 61. Sketches of A. transtensional strike-slip basins (in the lower part: transtensional La Gonzales Basin (NW-ern Venezuelan Andes) in a prominent releasing bend of the Bocono Fault -late Tertiary), and B. traspressional basins (modified, after Busby and Ingersoll 1995). The development of MHT is highly controlled by strike-slip faults as described previously. The re-activation of the major strike-slip faults in the MHT during alternating stress regimes is in accordance with Reading s Cycle, which predicts that every strike-slip fault may undergo alternating periods of extension and compression, while the slip directions adjust along major crustal faults (e.g. Reijs and McClay 1998, Barnes et al. 2001). Dextral strike-slip reverse faults along NW-SE orientation under a transpressional regime was followed by sinistral strike-slip faults along the same orientation related to a compressional event of different dynamics, while extensional 144

173 tectonics with low- and high-angle normal faulting controlled the subsidence of the last stages of MHT. The strike-slip faults developed obliquely across the trough during the Middle-Upper Miocene (T5 event) is characteristic for mature stages of a strike-slip basin s evolution (e.g., Doodley and McClay 1997). The compressional and flower structures occurring in places (e.g., Spileo, Theotokos areas) and the inferred transpressional regime for the first two stages of the MHT evolution, show that transpressional basins must have developed in places during Eocene Oligocene periods. A. B. Fig. 62. A. Strike-slip/ wrench basins (Einsele 2000). Transform motions may be associated either with a tensional component (transtensional) or a compressional component (transpressional). B. Sketches illustrating the Ridge Basin of California (after Crowell 1982). Left: The Basin s formation at a sigmoid bend of Saint Andreas fault, showing how the curvature of a strike-slip fault may produce an extensional basin closely adjacent to compressional uplift with superimposed tectonic pattern (based on Kingma 1958, Wilcox, Harding and Seely 1973, Crowell 1974a). Right: The basin s tectonic development. 145

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