A. SEGHEDI

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Palaeozoic Formations from Dobrogea and

Pre-Dobrogea – An Overview

ANTONETA SEGHEDI

National Institute of Marine Geology and Geoecology, 23−25 Dimitrie Onciul Street,

024053 Bucharest, Romania (E-mail: seghedi@geoecomar.ro)

Received 19 January 2011; revised typescript received 02 November 2011; accepted 11 December 2011

Abstract: An overview of lithological, palaeontological and geochronological evidence existing for the Palaeozoic

formations from Dobrogea and Pre-Dobrogea has enabled a better understanding of the Palaeozoic history of these

areas. Th e Lower Palaeozoic of Pre-Dobrogea, in places in continuity with the pelitic-silty facies of the underlying

Vendian (Ediacaran) deposits, was one of the peri-Tornquist basins of Baltica, suggesting that the Scythian Platform

in the Pre-Dobrogea basement represents the rift ed margin of the East European Craton. In North Dobrogea two

types of Palaeozoic succession have formed in diff erent tectonic settings. Deep marine Ordovician–Devonian deposits,

including pelagic cherts and shales, associated with turbidites, and facing Devonian carbonate platform deposits of the

East European Craton, form northward-younging tectonic units of an accretionary wedge, tectonically accreted above

a south-dipping subduction zone. South of the accretionary prism, the basinal to shallow marine Silurian–Devonian

deposits of North Dobrogea, showing a similar lithology to the East Moesian successions, accumulated on top of low-

grade Cambrian clastics with Avalonian affi nity indicated by detrital zircons. Late Palaeozoic erosion was accompanied

by deposition of continental alluvial, fl uvial and volcano-sedimentary successions, overlying their basement above

an imprecise Carboniferous gap. Th e low-grade metamorphic Boclugea terrane, showing Avalonian affi nity, and the

associated Lower Palaeozoic deposits represent East Moesian successions, docked to Baltica by the Lower Devonian

and subsequently involved in the Hercynian orogeny, being aff ected by Late Carboniferous–Early Permian regional

metamorphism and granite intrusion. Th e Late Carboniferous–Early Permian syn-tectonic sedimentation, regional

metamorphism of Palaeozoic formations and development of a calc-alkaline volcano-plutonic arc indicate an active

plate margin setting and an upper plate position of the Măcin-type successions during the Variscan collision, when the

Orliga terrane, with Cadomian affi nity, was accreted to Laurussia along a north-dipping subduction zone of the Rheic

Ocean. Th e East Moesian Lower Palaeozoic succession, overstepping its Ediacaran basement, represents an Avalonian

terrane, docked to the Baltica margin in the Early Palaeozoic. A narrow terrane detached from the Trans-European

Suture Zone (TESZ) margin of the Baltica palaeocontinent forms a tectonic wedge within the East Moesian basement.

Th e Palaeozoic sedimentary record of East Moesia shows a quartzitic facies in the Ordovician, graptolite shales in

Upper Ordovician–Wenlock, black argillites in the Ludlow-Pridoli and fi ne-grained clastics in the Lower Devonian.

Eifelian continental sandstones are followed by a carbonate platform from Givetian to Tournaisian times and coal-

bearing clastics in the Carboniferous, indicating a foredeep basin evolution. By the Eifelian both East Moesia and

Pre-Dobrogea were part of Laurussia, sharing the same old red sandstone facies. Th e Permian is a time of rift ing in

Dobrogea and Pre-Dobrogea, although evidence for rift ing in the East Moesian sedimentary record is very limited.

In the eastern basins of Pre-Dobrogea, Permian rift ing was accompanied by alkaline bimodal volcanism of the basalt-

trachyte association, that aff ected also the northern margin of North Dobrogea. Late Permian within-plate alkaline

magmatic activity emplaced plutonic and hypabyssal complexes along the south-western margin of North Dobrogea.

Th e model proposed for the Palaeozoic history based on existing data for the north-western margin of the Black Sea

records early Palaeozoic docking to Baltica of the Avalonian terrane of East Moesia, including the Boclugea terrane of

North Dobrogea. Late Carboniferous–Early Permian accretion of the Cadomian Orliga terrane from North Dobrogea,

accompanied by Hercynian metamorphism and granite intrusion, correlates with the closure of the Rheic Ocean.

Subsequently, Avalonian and Cadomian terranes, together with a narrow terrane detached from the TESZ margin of

Baltica palaeocontinent, were displaced southward along the strike-slip fault system of the TESZ.

Key Words: North Dobrogea Orogen, Moesian Platform, Scythian Platform, lithology, biostratigraphy

Dobruca ve Ön-Dobruca’nın Paleozoyik Formasyonları

Özet: Dobruca ve Ön-Dobruca’daki Paleozoyik formasyonlarının litolojik, paleontolojik ve jeokronolojik özelliklerinin

gözden geçirilmesi bu bölgelerin Paleozoyik tarihçelerinin daha iyi anlaşılmasını sağlar. Ön-Dobruca’nın Alt Paleozoyik

Turkish Journal of Earth Sciences (Turkish J. Earth Sci.), Vol. 21, 2012, pp. 669–721. Copyright ©TÜBİTAK

doi:10.3906/yer-1101-20 First published online 11 December 2011

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670

Introduction

Th e western margin of the East European Craton, a

major terrane boundary along the contact between the

stable Precambrian Fennoscandian-East European

Craton and the younger structures of Western and

Southern Europe, was defi ned as the Trans-European

Suture Zone or TESZ (Pharaoh 1999) (Figure 1).

Along the TESZ, peri-Gondwanan terranes of Far

East Avalonia are found, accreted to the former

Baltica palaeocontinent during the Lower Palaeozoic

(Ziegler 1986, 1988; Pharaoh 1999; Winchester et

al. 2002, 2006), are mingled with proximal Baltican

terranes. All these terranes with Avalonian and

Baltican affi nity are variously displaced together

along the strike-slip faults of the TESZ (Winchester

et al. 2002, 2006; Nawrocki & Poprawa 2006; Oczlon

et al. 2007). Th e southeastern segment of the TESZ

runs through the southwestern part of Ukraine

and Moldavia, continuing to the Black Sea through

south-eastern Romania.

Th e northwestern margin of the Black Sea includes

three major structural units: the westernmost segment

of the Scythian Platform, the North Dobrogea Orogen

and the Moesian Platform. All these units include

Palaeozoic formations, concealed in the platforms

and exposed in North Dobrogea. In order to better

understand the geological evolution of this area and

improve palaeogeographic models, it is important to

establish or update terrane affi nities. Due to limited

reliable information and still poorly defi ned terrane

affi nities, most palaeocontinental reconstructions

for Moesia and/or Dobrogea assume that each forms

one single terrane (Mosar & Seghedi 1999; Stampfl i

2000; Kalvoda et al. 2002; Cocks & Torsvik 2005,

istifl eri, Ön-Dobruca’nın bazı bölgelerinde daha altta yer alan Vediyen (Edikariyen) pelitik-siltli fasiyeslerle devamlılık

gösterir, ve Baltika’nın peri-Tornquist havzalarından birini oluşturur. Bu durum Ön-Dobrucayı da içine alan İskit

Platformu’nun Doğu Avrupa Kratonu’nundan rift leşme ile ayrıldığına işaret etmektedir. Kuzey Dobruca’da farklı

tektonik ortamlara işaret eden iki tip Paleozoyik istif bulunur. Çört ve şeyl ve bunlarla ilişkili türbiditlerden oluşan

derin denizel Ordovisyen–Devoniyen çökelleri, ve bu çökellerin Doğu Avrupa Kratonuna bakan kenarında gelişmiş

Devoniyen karbonat platformu, güneye doğru dalan bir dalma-batma zonunda gelişmiş bir eklenir prizma oluşturur.

Eklenir prizmanın güneyinde, derinden sığ denize kadar değişen Kuzey Dobruca’nın Siluriyen–Devoniyen çökelleri,

Doğu Moezya istfl erine benzerlik gösterir, ve kırıntılı zirkonlarla Avalonya’ya bağlı olduğu saptanan düşük dereceli

Kambriyen kırıntılıları üzerinde çökelmiştir. Geç Paleozoyik’te gelişen erozyon ve bölgenin karalaşması sonucu karasal

çökeller ve volkanik kayaları, arada bir Karbonifer boşluğu olmak üzere bu temel üzerinde yer alır. Düşük dereceli

metamorfik kayalardan oluşan Bokluca mıntıkası Avalonya özellikleri gösterir, ve Bokluca’ya bağlı Alt Paleozoyik

kayaları Doğu Moezya özellikleri taşır; bu birimler Erken Devoniyen’de Baltika ile çarpışmış ve daha sonra Geç

Karbonifer–Erken Permiyen’de rejyonal metamorfizma ve granitik sokulumlar ile tanımlanan Hersiniyen orojenezi

geçirmiştir. Bu özellikler ve Geç Karbonifer–Erken Permiyen yaşlı tektonizma ile eşyaşlı sedimentasyon, bölgenin bu

dönemde aktif bir kıta kenarı konumunda olduğuna işaret eder. Kadomiyen özellikler gösteren Orliga mıntıkası, Reik

Okyansu’nun kuzeye doğru dalıp yok olması sonucu Lavrasya’ya eklenmiştir. Doğu Moezya’nın Alt Paleozoyik istifl eri,

Erken Paleozoyik’te Baltika’ya yamanan Avalonya tipi bir mıntıkaya aittir. Baltika’nın Trans-Avrupa Kenet Zonu (TESZ)

kıta kenarınan ayrılmış ince bir mıntıka Doğu Moezya temeli içinde bir kıymık oluşturur. Doğu Moezya’nın sedimenter

istifi, Ordovisyen’de kuvarsitik fasiyesler, Üst Ordovisyen–Venlok’ta graptolitli şeyller, Ludlov–Pridoli’de siyah çamur

taşları, Alt Devoniyen’de ince taneli kırıntılılardan yapılmıştır. Efyeliyen yaşlı karasal kumtaşlarını takiben Givetiyen–

Turnaziyen zaman aralığında platform karbonatları gelişmiş, ve daha sonra Karbonifer’de kömür içeren kırıntılılar, bir

ön-ülke havzasında çökelmiştir. Eyfeliyen’de hem Ön-Dobruca hem de Doğu Moezya, Lavrasya’nın yamanmıştır ve

benzer kırmızı kumtaşı fasiyesleri gösterirler. Permiyen’de Dobruca ve Ön-Dobruca’da rift leşme gözlenir, buna karşın

Doğu Moezya’da rifitleşme ile ilgili çökel kayıtları çok kıtdır. Ön-Dobruca’nın doğu havzalarında Permiyen rift leşmesi

ile beraber bazalt-trakit birlikteliğinden oluşan alkalin bimodal volkanizma gelişmiş, ve bu volkanizma Kuzey

Dobruca’nın kuzey kenarını da etkilemiştir. Geç Permiyen levha-içi alkalin magmatizma sonucu Kuzey Dobruca’nın

güneybatı kenarı boyunca derinlik ve yarı-derinlik kayaları yerleşmiştir. Burada sunulan model, Erken Paleozoyik’te

Baltika’nın güney sınırı boyunca Doğu Moezya ve Kuzey Dobruca’nın Bokluca mıntıkasını içeren Avalonya kökenli kıta

parçaçıklarının Baltika’ya eklenmesini içerir. Kuzey Dobruca’nın Kadomiyen kökenli Orliga mıntıkası Geç Karbonifer–

Erken Permiyen’de kuzeye eklenmiş ve bu olay sonucu Reik Okyanusu kapanarak Hersiniyen metamorfizması ve

granit yerleşimi gerçekleşmiştir. Bu olayları takiben Avalonya ve Kadomiyen kökenli mıntıkalar, ve Baltika’nın TESZ

kenarından kopan ince bir mıntıka, TESZ boyunca gelişen doğrultu-atımlı faylar boyunca güneye doğru ötelenmiştir.

Anahtar Sözcükler: Kuzey Dobruca orojeni, Moezya Platformu, İskit platformu, litoloji, biyostratigrafi

A. SEGHEDI

671

2006; Nawrocki & Poprawa 2006; Winchester et al.

2006). However, from detailed analysis of various

existing data, terranes with both Baltican and

Avalonian palaeogeographic affi nities were inferred

to make up the Moesian Platform and a model for

their displacement along the TESZ, together with

the North Dobrogea terrane, was proposed (Oczlon

et al. 2007). Detrital zircon data enabled separation

of Avalonian and Cadomian terranes in the North

Dobrogea metamorphic suites, brought together

following the closure of the Rheic Ocean (Balintoni

et al. 2010).

Th e goal of this paper is to provide an

overview of the lithological, biostratigraphical and

geochronological information from the Palaeozoic

formations in the northwest Black Sea area and

comment on the evidence for Palaeogeographic affi nities. Th e Palaeozoic record from Pre-Dobrogea presented here is the result of correlation of borehole data across state borders, based on the petrographic studies of thin sections provided by the former Oil Institute in Bucharest and the Geological Institute from Kishinev. Core samples stored at the Geological Institute from Kishinev and the Geological Institute of Romania have also been examined. For North Dobrogea, the review of geological data is based on papers published in local journals, abstracts and fi eld guide books, as well as on unpublished reports and PhD theses. Th e data on East Moesia are summarized according to the synthesis of the Moesian Palaeozoic from Romania (Seghedi et al. 2005a, b), to serve as a basis for comparison with North Dobrogea and Pre-Dobrogea.

Figure 1. Location of Dobrogea on a simplifi ed terrane map of Europe (modifi ed aft er the TESZ map of EUROPROBE project).

US– Ukrainian shield, VM– Voronezh massif, EC– East Carpathians, SC– South Carpathians, SP– Scythian Platform,

MP– Moesian Platform, NDO– North Dobrogea orogen.

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Geological and Tectonic Background

Th e north-western margin of the Black Sea is a highland area referred to as Dobrogea, a geographical and historical province confi ned between the Black Sea shore, the Sfântu Gheorghe Distributary of the Danube Delta and the lower reaches of the Danube River. Dobrogea is surrounded by the lowlands of the Pre-Dobrogea Depression in the north and the Romanian Plain in the west (Figure 2). Th e largest

part of Dobrogea belongs to Romania, except for its southern margin that continues for some distance in Bulgaria. North of the Dobrogea highlands, the fl at-lying Pre-Dobrogea Depression is shared by three countries: Romania, Moldavia and Ukraine.

Th e area consists of three main tectonic units, two Palaeozoic platforms (Moesian and Scythian) and the Cimmerian Orogen of North Dobrogea (Săndulescu 1984) (Figure 3). While the Scythian Platform is

Figure 2. Th e main geotectonic units of the western Black Sea margin. Th e area with darker grey shading represents

exposures of pre-Cenozoic rocks. EM– East Moesia; WM– West Moesia; SGF– Sfantu Gheorghe Fault; PCF–

Peceneaga-Camena Fault; COF– Capidava-Ovidiu Fault; PF– Palazu Fault; EF– Eforie Fault; IMF– Intra-

Moesian Fault.

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673

entirely concealed by Quaternary deposits, the

eastern parts of the North Dobrogea orogen and of

the Moesian Platform are exposed in North Dobrogea

and Central and South Dobrogea, respectively.

Separated from the Scythian Platform by the

Sfântu Gheorghe Fault and bounded southward

by the Peceneaga-Camena Fault, North Dobrogea

represents the south-eastern part of the North

Dobrogea Cimmerian Orogen, where the Hercynian

basement and its Mesozoic cover are exposed (Figures

2 & 3). Th e north-western part of the belt is covered

by Cenozoic deposits of the Carpathian foredeep.

Th e stratigraphic gap between the Late Jurassic

and Cenomanian sediments which seal both the

Cimmerian structures of North Dobrogea, as well

as the Peceneaga-Camena Fault, suggests that

throughout the Early Cretaceous North Dobrogea

was part of the northern rift shoulder of the Western

Black Sea Basin, which sourced the kaolinite-rich

Aptian syn-rift sediments preserved both in the Pre-

Dobrogea depression and South Dobrogea (Rădan

1989; Ion et al. 2002).

Surrounded by the Carpathians and the Balkans,

the Moesian Platform is an area with a heterogeneous

and complex Precambrian basement overlain by a

thick Palaeozoic to Cenozoic cover. West of the Black

Sea, the eastern part of the platform (East Moesia)

consists of two tectonic provinces separated by the

Capidava-Ovidiu Fault (Figure 3).

Confi ned between the Peceneaga-Camena and

Capidava-Ovidiu crustal faults, Central Dobrogea

exposes the Neoproterozoic Moesian basement.

Th e fl at-lying Palaeozoic cover, preserved only

west of the Danube, above the subsided part of the

Figure 3. Schematic map showing the main structural units of Dobrogea (modifi ed from Seghedi et al. 2005a). ND– North Dobrogea;

CD– Central Dobrogea; SD– South Dobrogea; LCF– Luncaviţa-Consul Fault; other abbreviations as in Figure 2.

Post-tectonic cover(Babadağ Basin)

Palaeozoic

platform cover

platform cover

PALAEOZOIC FROM DOBROGEA AND PRE-DOBROGEA

674

Neoproterozoic basement, has been completely

removed from the uplift ed block of Central Dobrogea

during an unspecifi ed period of pre-Bathonian

erosion. In the outcrop area, the Neoproterozoic

basement is unconformably covered by Bathonian–

Kimmeridgian carbonate platform successions above

local remnants of a pre-Bathonian weathering crust

(Rădan 1999) (Figure 3).

South Dobrogea is a subsided Moesian block,

along the Capidava-Ovidiu and Intramoesian faults.

Th is block exposes only the Mesozoic–Cenozoic

Moesian cover with frequent discontinuities and

gaps. West and north-west of the Danube River, the

corresponding parts of these main units of Dobrogea,

concealed by Cenozoic deposits, lie at depths of over

600 m in the footwall of the Danube Fault (Gavăt et

al. 1967) (Figure 2).

Th e Pre-Dobrogea Depression

Th e Pre-Dobrogea depression represents a Mesozoic–

Tertiary depression superimposed on a pre-Triassic

basement. According to Săndulescu (1984), the Pre-

Dobrogea basement is the westernmost segment of

the epi-Variscan Scythian Platform. Running E–W

along the south-western corner of the East European

Craton, the Scythian Platform is buried westward

beneath the Tertiary molasse foredeep of the East

Carpathians.

Both the age of cratonization and Variscan

history of this tectonic unit are quite controversial.

Located between the East European Craton and

the Alpine-Cimmerian folded belts on its southern

border, the Scythian Platform is classically defi ned

as a wide Variscan belt referred to as the ‘Scythian

orogen’ (Mouratov & Tseisler 1982; Milanovsky 1987;

Zonenshain et al. 1990; Nikishin et al. 1996). With

active orogenesis supposed to have occurred from

Early Carboniferous to Permian times (Nikishin et

al. 1996, 2001; Stampfl i & Borel 2002), it has usually

been considered to represent the link between the

Variscan orogen of western and central Europe

and the Uralian belt at the eastern edge of the East

European Platform. In the Mesozoic, the ‘Scythian

orogen’ showed a platform stage of development

(Mouratov 1979), its basement being concealed by

the superimposed Mesozoic–Tertiary deposits of the

Pre-Dobrogea Depression. Other authors regard this area as the southern passive margin margin of the East European craton reworked by Late Proterozoic (Early Baikalian) and younger tectonism (Kruglov & Tsypko 1988; Gerasimov 1994; Drumea et al. 1996; Milanovsky 1996; Pavliuk et al. 1998; Poluchtovich et al. 1998; Stephenson 2004; Stephenson et al. 2004; Saintot et al. 2006).

Th e basement of the Pre-Dobrogea depression is a highly tectonized area about 100 km wide, defi ned against the neighbouring units by the crustal faults Baimaklia-Artiz (or Leovo-Comrat-Dnestr) and Sfântu Gheorghe, known from geophysical and drilling data (Figure 4). Th e structure of the pre-Mesozoic basement is defi ned by the intersection of two major fault systems. A system of WNW–ESE-trending, parallel faults has controlled the development of intrabasinal longitudinal ridges. Th is is intersected by a N–S-trending fault system, best developed east of the Prut River (Neaga & Moroz 1987; Drumea et al. 1996; Ioane et al. 1996; Visarion & Neaga 1997).

Beneath the fl at-lying Middle Jurassic –Tertiary cover, the area of the Pre-Dobrogea Depression is a Permian palaeorift (Neaga & Moroz 1987). A longitudinal high, the Bolgrad-Chilia (or Lower Prut) Horst (Neaga & Moroz 1987; Moroz et al. 1997; Visarion & Neaga 1997), separates two depressions elongated E–W (Figures 4 & 5). Th e northern depression includes the Sărata-Tuzla and Aluat basins, separated by the Orehovka-Suvorovo basement high. Th e Sărata-Tuzla basin is complicated by a minor longitudinal intrabasinal ridge. Th e Aluat basin continues into Romania in the WNW-elongated Bârlad depression; this basin shows a staircase geometry, its bottom being progressively downthrown westward towards the East Carpathian foredeep. Th is is accommodated by a system of N–S faults, reactivated as result of nappe stacking in the East Carpathians. Th e Sulina (or Lower Danube) basin, with its depocentre situated in the Danube Delta, developed south of the Bolgrad-Chilia high (Figure 4).

Th e basement lies at depths of 1–1.5 km in areas of basement highs and at depths of 3–4 km in depressions (Mouratov & Tseisler1982). Th e basement of the Pre-Dobrogea consists of magmatic

A. SEGHEDI

675

rocks (granites, diorites and gabbros), that yield

Neoproterozoic K-Ar ages (790, 640–620 Ma) (Neaga

& Moroz 1987) (Figure 5). In the central part of the

Bolgrad-Chilia and Orehovka-Suvorovo highs, the

magmatic basement is unconformably overlain

by undeformed Vendian deposits. Th ese deposits

were intersected by the Orehovka and Suvorovo

boreholes for a thickness of over 2000 m. Remnants

of a palaeo-weathering crust were found on top

of the magmatic basement rocks (Neaga & Moroz

1987). Th e Neoproterozoic Ediacaran (Vendian) age

is derived from a phytocenosis with Vendotaenia

antiqua Grujilov, identifi ed in the Orehovka borehole

(Visarion et al. 1993). A K-Ar age of 600±20 Ma yielded

by pelitic rocks is consistent with the palynological

data. Th e Vendian succession (Avdărma Series) is

upward shallowing, including marine sediments

(conglomerates, sandstones, volcanic sands, black

Figure 4. Distribution of the main basins of the Scythian Platform in the Predobrogea depression, with location of the main boreholes

(compiled aft er Neaga & Moroz 1987; Paraschiv 1986; Pană 1997). Shades of grey represent grabens and highs, respectively.

PALAEOZOIC FROM DOBROGEA AND PRE-DOBROGEA

676

shales rich in phosphate nodules and grey siltstones, sandstones and mudstones) grading upward into continental deposits (conglomerates and purple

pebbly sandstones, with siltstone and mudstone interbeds). With these features, the Vendian deposits of the Pre-Dobrogea resemble the Avdărma Series

Figure 5. Mesozoic subcrop map of the Scythian Platform, showing the distribution and structure of the Neoproterozoic and

Palaeozoic formations (compiled aft er Visarion et al. 1993; Ioane et al. 1996 and Visarion & Neaga 1997).

A. SEGHEDI

677

from the East European Platform cover, exposed

along the Dnestr River, the only diff erence being the

lesser thickness of the latter (only 600–700 m).

Th e platform cover of the Pre-Dobrogea comprises

mainly Cretaceous to Quaternary sediments with

a total thickness of 2500 to 3500 m (Permyakov &

Maidanovich 1984). Due to the high mobility of the

area, repeatedly subjected to oscillatory movements,

the sedimentary cover shows stratigraphic gaps and

unconformities, being characterized by the absence

of the Lower Jurassic and thin Lower Cretaceous

deposits (Ionesi 1994). According to the synthesis

of Ion et al. (2002), the Mesozoic cover includes

Callovian–Oxfordian black shales in the Danube

Delta and carbonates in the rest of the Pre-Dobrogea,

overlain by Oxfordian–Kimmeridgian carbonate

platform limestone. Dolomites, clastics and

evaporites develop in the Berriasian–Valanginian;

dolomites, dolomites and clastics in the Middle–late

Aptian, while clastics occupy the northern half of

the Delta in the Sarmatian followed by Late Meotian

shales, and the Quaternary is represented by sand

and clay of marine, fl uvial and lacustrine origin.

North Dobrogea

Th e narrow, NW-trending Cimmerian fold and

thrust belt of the North Dobrogea orogen (Figure 3)

is a basement with a Hercynian history of magmatism

and deformation. Th is basement was subsequently

involved in early Alpine (Cimmerian) events,

experiencing extension during the Late Permian–

Middle Triassic, and compression (transpression) in

the Late Triassic–Middle Jurassic (Seghedi 2001). A

major high-angle reverse fault with a NW–SE strike

(Luncaviţa-Consul Fault, Savul 1935), represents

the contact between the Măcin and Tulcea zones

(Mutihac 1964) (Figure 6). Th ese zones refer to

areas with dominantly Palaeozoic and Triassic

exposures respectively; Jurassic deposits occur in

limited areas in both zones. Th e Măcin zone (Măcin

Nappe, Săndulescu 1984) is a Cimmerian tectonic

unit exposing mainly Palaeozoic formations in the

western part of North Dobrogea. Th e Tulcea zone is

the larger, eastern part of North Dobrogea, exposing

mostly Triassic and Jurassic formations included in

three Cimmerian thrusts (Săndulescu 1984).

All four major Cimmerian thrust bounded units

recognised in North Dobrogea have Hercynian

deformed basement and Triassic or Triassic–Jurassic

cover (Mirăuţă, in Patrulius et al. 1973; Săndulescu

1984) (Figure 7). Th e Cimmerian tectonic units are

interpreted either as low-angle nappes, based largely

on geophysical data (Săndulescu 1984; Visarion et al.

1993), or as high-angle thrusts, based on borehole

information (Baltres 1993). Th e latter tectonic model

is fi gured in the transect VII of the TRANSMED

Atlas (Papanikolaou et al. 2004). Th e Măcin

Cimmerian tectonic unit (corresponding to the

descriptive Măcin zone), exposes largely Palaeozoic

formations, with only minor exposures of Triassic

deposits (Figures 3 & 6). In the Tulcea zone, exposing

mainly Triassic deposits and a few Palaeozoic and

Jurassic formations, are three Cimmerian tectonic

units (Săndulescu 1984): the Consul, Niculiţel and

Tulcea nappes.

Th e Triassic succession, unconformable on the

Hercynian basement, starts with lower Scythian

(Werfenian) continental fanglomerates, followed

by sandstones and upper Scythian limestone

turbidites (Baltres 1993). Rhyolites and basalts

started to be emplaced in the late Scythian, with

the basaltic volcanism continuing up to Middle

Anisian (Baltres et al. 1992). Th e basaltic volcanism

is partly coeval with deposition of nodular and

bioturbated limestones and cherty limestones. Th e

basinal succession terminates with Late Anisian–

late Carnian Halobia marls (Baltres et al. 1988).

Turbiditic deposits accumulated, starting in the late

Carnian and ranging up to the middle Jurassic, with

an upward coarsening trend of coarse members

(Baltres 1993; Grădinaru 1984). Shallow marine

carbonate sedimentation took place along the basin

margin from the Scythian to the Norian and in the

Oxfordian–Kimmeridgian (Grădinaru 1981; Baltres

1993). Apart from numerous unpublished reports, a

detailed presentation of the Triassic–Jurassic deposits

of North Dobrogea is given in Grădinaru (1981, 1984,

1988) and Seghedi (2001).

An early Cretaceous history is not preserved in

the stratigraphic record, except for scarce remnants

of a kaolinitic weathering crust, found in several

locations on top of the the Hercynian basement

or the Triassic deposits. In one locality this crust

PALAEOZOIC FROM DOBROGEA AND PRE-DOBROGEA

678

is preserved beneath transgressive Cenomanian

calcarenites and was ascribed to the Aptian (Rădan

1989). Th e post-tectonic cover of the North Dobrogea

orogen is represented by the shallow-marine Upper

Cretaceous sediments of the ‘Babadag basin’. Th ey

seal the deeply truncated Hercynian and Cimmerian

structures, overstepping the easternmost segment

of the Peceneaga-Camena Fault and overlying the

Histria Formation (Figure 3). Th e Upper Cretaceous

succession includes Albian bioclastic limestones,

Cenomanian–Turonian detrital limestones with

conglomerate interbeds in the eastern part, Coniacian

detrital limestones with cherts, chalk and glauconite

and Santonian–Campanian nodular limestones

developed only along the eastern margin of North

Dobrogea (Ion et al. 2002 and references therein).

Figure 6. Quaternary subcrop map showing the outcrop areas of Palaeozoic formations and the main Cimmerian faults in

North Dobrogea (modifi ed from Seghedi 1999).

A. SEGHEDI

679

Fig

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PALAEOZOIC FROM DOBROGEA AND PRE-DOBROGEA

680

Central Dobrogea

Th e Central Dobrogea basement consists of two terranes with distinct lithologies and deformational history. Metapelites and metabasites of the Altin Tepe Group, with an initial amphibolite facies metamorphism, are ascribed to the Late Proterozoic based on K-Ar ages on biotite (696–643 Ma, Giuşcă et al. 1967) (Figure 8). Metabasites are tholeiitic, showing arc/back arc affi nities (Crowley et al. 2000). Th e mesometamorphic rocks are exposed in an antiform south of the Peceneaga-Camena Fault and develop a wide greenschist facies mylonitic zone (Mureşan 1971) along their contact with the overlying Histria Formation. Th is ductile shear zone in Altin Tepe rocks contrasts with the brittle deformation in the Histria Formation lithologies, the contact showing the main characteristics of a low-angle detachment fault (Seghedi et al. 1999).

Well exposed over the entire Central Dobrogea area, the Histria Formation is a succession up to 5000 m thick (O. Mirăuţă 1965, 1969; Visarion et al. 1988), consisting of two coarse members of sandstone dominated, channelized midfan turbidites, separated by a thinner member (up to 500 m thick) of distal, abyssal plain turbidites (Seghedi & Oaie 1995; Oaie

1999). Palaeofl ow directions indicate a southern

source area which supplied both terrigenous and

volcanic clasts (Oaie 1999). Th e composition of

coarse members suggests that a major continental

margin source delivered Palaeoproterozoic BIF and

gneisses into the basin; a second, volcanic source,

yielded basalt and rhyolite clasts (Oaie et al. 2005).

Detrital zircon ages are consistent with Archaean and

Palaeoproterozoic sources (Żelaźniewicz et al. 2009;

Balintoni et al. 2011). Sedimentological, structural

and mineralogical data suggest that the Histria

Formation accumulated in a foreland basin setting

(Seghedi & Oaie 1995; Oaie 1999; Oaie et al. 2005), an

interpretation consistent with results of geochemical

and detrital zircon distribution data (Żelaźniewicz et

al. 2009).

Th e Late Proterozoic–Early Cambrian

depositional age of the turbidites is constrained

by palynological associations (Iliescu & Mutihac

1965). A medusoid imprint in the middle member

of the Histria Formation, identifi ed as Nemiana

simplex Palij, suggests an Ediacaria-type fauna (Oaie

1992). In boreholes from the Romanian Plain, the

turbidites are unconformably overlain by fl at-lying

quartzitic sandstones with Ordovician graptolites (E.

Figure 8. Schematic tectonostratigraphic charts for the metamorphic rocks of Dobrogea. K-Ar ages are taken from Ianovici & Giuşcă

(1961); Giuşcă et al. (1967); Kräutner et al. (1988); Ar-Ar ages from Seghedi et al. (1999); monazite CHIME ages from Seghedi

et al. (2003a); U-Pb detrital zircon ages from Balintoni et al. (2010, 2011).

A. SEGHEDI

681

Mirăuţă 1967; Iordan 1992, 1999). Neoproterozoic

deformation of the turbiditic succession in very-

low-grade metamorphic conditions occurred at 572

Ma (K-Ar WR) (Giuşcă et al. 1967) and resulted in

E–W-trending, open normal folds, with axial-planar

slaty cleavages, penetrative only in fi ne-grained

lithologies (O. Mirăuţă 1969; Seghedi & Oaie 1995).

Detrital zircon ages (Żelaźniewicz et al. 2001) are

interpreted to indicate an Avalonian affi nity for the

Ediacaran turbidites (Oczlon et al. 2007) and a peri-

Amazonian provenance is suggested, based on the

age distribution pattern (Balintoni et al. 2011).

Th e platform cover exposed in Central Dobrogea

starts with Bathonian calcarenites rich in crinoidal

debris, transgressive on the Ediacaran basement,

followed by Callovian–Oxfordian carbonates showing

the same facies as in the Pre-Dobrogea depression

and overlain by Oxfordian–Kimmeridgian carbonate

platform limestones (Ion et al. 2002 and references

therein). Only in the northeast does the Upper

Cretaceous cover of North Dobrogea overstep the

Ediacaran basement (Figure 3).

South Dobrogea

In the subsurface of South Dobrogea, the cratonic

basement of the Moesian Platform is comparable to

that of the Ukrainian Shield, containing Archaean

gneisses and an Lower Proterozoic banded iron

formation (BIF) (Palazu Mare Group) (Giuşcă et al.

1967, 1976; Giuşcă 1977) (Figure 8). Th e gneisses

and the banded iron formation, correlated with

the Ukrainian shield of the East European Craton

(Giuşcă et al. 1967; Kräutner et al. 1988), have been

interpreted as the small, proximal Baltican Palazu

terrane, displaced along the TESZ (Oczlon et al.

2007).

Th e BIF shows a Svecofennian HT-LP amphibolite

facies metamorphism, with andalusite-sillimanite

assemblages (Giuşcă et al. 1967, 1976). A later,

greenschist-facies retrogression was correlated

with the very low-grade metamorphism of the

late Cadomian Neoproterozoic cover (Kräutner

et al. 1988). Neoproterozoic volcano-sedimentary

deposits (Cocoşu Formation) include volcanics and

volcano-sedimentary successions derived from a

mafi c, alkaline volcanism (Figure 8). Basaltic fl ows

of basanites and trachybasalts showing intraplate geochemical affi nities resulted through the rift ing of the cratonic basement (Seghedi et al. 2000). Th e Late Proterozoic deformation of the basaltic rocks, dated at 547 Ma (K-Ar WR), is assumed to be connected to northward thrusting of Archean and Palaeoproterozoic suites (Giuşcă et al. 1967; Kräutner et al. 1988).

According to the geological record, the undeformed sedimentary cover of South Dobrogea includes Ordovician to Quaternary formations separated by gaps. Several cycles have been separated: Cambrian–Westphalian, Permian–Triassic, Bathonian–Eocene and Miocene–Quaternary (Paraschiv 1975; Ionesi 1994). Th e Palaeozoic formations will be described in the next chapter.

Th e Mesozoic platform cover starts with unconformable, scarce Triassic red-beds, followed by Middle–Upper Triassic calcareous successions. Th e Jurassic includes mainly carbonate platform sediments; strongly dolomitized limestones and calcarenites prevail, unlike in the Central Dobrogea. Along the Capidava-Ovidiu Fault, the typical marine patch-reef facies of South Dobrogea is replaced by regressive deposits, with alternating marine limy and evaporitic-lagoonal facies, partly coeval with those of the Oxfordian–Tithonian and Kimmeridgian–Tithonian respectively (Avram et al. 1993).

Twelve distinct Cretaceous formations separated by stratigraphic discontinuities and established through detailed biostratigraphic studies on outcrops and boreholes in South Dobrogea are presented in several papers (Avram et al. 1993; Ion et al. 2002 and references therein). Th e Berriasian–Valangianin–Early Hauterivian includes an evaporitic-detrital succession; limestones with clayey interbeds, calcarenitic, dolomite-clayey, coarse siliciclastic, calcareous-detrital and calcareous-marly successions. Th e Middle–Upper Aptian has a fl uvio-lacustrine facies with red beds, coal and kaolinite, while in the Upper Aptian–Albian onshore clastics are followed by a marine marly-silty facies. Th e transgressive Lower Cenomanian consists of a basal conglomerate, glauconitic chalk and upper, massive chalky sandstone. Th e marl-dominated Upper Cenomanian is overlain by clastic Turonian–Santonian–Campanian chalky-glauconitic-quartzitic sandstone

PALAEOZOIC FROM DOBROGEA AND PRE-DOBROGEA

682

with inoceramus shell-debris, followed by thick chalk with chert. Th e Lower Maastrichtian includes a lacustrine, variegated clayey and marly succession, followed by Upper Maastrichtian interbedded chalky marls and clays, chalky glauconitic sands/sandstones and massive chalky limestones. Th e Palaeogene includes glauconitic sands and sandy biocalcarenites, rich in nummulites. Th e Miocene succession consists of normal marine Upper Badenian (Kossovian) and brackish Sarmatian clastics and bioclastic limestones, separated by a break in sedimentation. Pliocene deposits include sands, conglomerates and silty clays, containing a mollusc fauna typical for the Upper Pontian, Dacian and Romanian. Th e Quaternary includes sands, clays and loess deposits.

Description of the Palaeozoic Formations

Th e Palaeozoic Record of Pre-Dobrogea Depression

Cambrian–Silurian – Th e areal distribution of

Palaeozoic lithologies in Pre-Dobrogea is shown on

the pre-Triassic subcrop map in Figure 5 and the

vertical succession of facies in Figure 9. A sequence of

reddish sandstones, siltstones and mudstones, 100–

300 m thick, intersected by boreholes in the Sărata-

Tuzla grabens was ascribed to the Cambrian (Bogatzev

1971, in Belov et al. 1987), or to the early Cambrian–

Silurian interval (Belov et al. 1987). Orthoquartzitic

sandstones from the Buciumeni borehole in Romania

(Figure 4) yielded a palynological association of

spores, acritarchs, leiosphaerids and algae specifi c to

the Lower Ordovician, possibly including the Lower

Cambrian (Paraschiv 1986a).

Th e presence of the Ordovician and Lower–

Middle Silurian was identifi ed, based on Chitinozoan

assemblages in Liman borehole 1 from the southern

margin of the Pre-Dobrogea Depression (Vaida &

Seghedi 1997). Th e deposits, penetrated to a thickness

of 2674 m, had been previously assigned entirely to

the Vendian. Th e lithofacies includes grey, bluish and

purple siltstones and mudstones, with sandstones

interbedded in the middle part and limestones in the

upper parts of the succession. Only the upper part

of the deposits (620 m thick) yielded palynological

associations dominated by chitinozoans, with

subordinate acritarcha and scolecodonts. Along

with the long range chitinozoans, there are species

that terminate their evolution in the Ordovician

(Conochitina brevis T. et de J., Rhabdochitina gracilis

Eis., Desmochitina urceolata B. et T., D. pellucida

B. et J., Eremochitina sp., E. baculata B. et de J.,

Siphonochitina sp. and acritarchs Baltisphaeridium

digitiforme Gorka, Lophosphaeridium papulatum

Martin, Multiplicisphaeridium continuatum

Kjellstrom.) and Silurian assemblages (Conochitina

gordonensis Cramer, C. intermedia Eis., Rhabdochitina

conocephala Eis, Desmochitina minor Eis., D. tinae

Cramer, the latter characteristic for the Llandovery)

(Vaida & Seghedi 1997). Th e rest of the succession

(1954 m) belongs to the Vendian–Cambrian. Th e

situation in the Liman borehole suggests that the

Ordovician–Middle Silurian deposits may be present

in other areas of the Pre-Dobrogea Depression.

Lower Devonian – Th e Lower Devonian (known

east of the Prut River as the Iargara Series) includes

fi ne-grained clastics showing the same facies as the

coeval sediments from the western part of the East

European Platform (Neaga & Moroz 1987). Th e

maximum thickness of deposits is 1200 m, attained in

the Kazaklia borehole. Th e Iargara Series represents

a dominantly terrigenous, upward coarsening and

shallowing sequence. Its lower part (Cociulia Beds)

consists of marine shales and argillites with thin

limestone interbeds, with a fauna of pelecypods,

brachiopods and Orthoceratid fragments. Th e middle

part (Lărguţa Beds) includes grey-greenish argillites,

interbedded with sandstones, siltstones, detrital and

bioclastic limestones, the latter rich in brachiopods,

pelecypods, bryozoans, ostracods and tentaculites

(Safarov & Kaptzan 1967). Th e upper part (Enichioi

Beds) comprises continental deposits, including

quartzitic sandstones with thin red shale interbeds,

ascribed to the Lower Devonian on stratigraphic

criteria. West of the Prut River, purple quartzitic

sandstones with Dictyonema sp., Umbellina bella and

Dentalina iregularis have been ascribed to the Lower

Devonian (Baltes 1969). Tuff s of andesitic, andesitic-

dacitic and dacitic composition are conformably

interbedded with the sandstones (Moroz et al. 1997).

Middle Devonian–Lower Carboniferous – Th e Middle–Upper Devonian includes evaporate-bearing,

A. SEGHEDI

683

Figure 9. Stratigraphic chart for the Palaeozoic deposits of the Scythian Platform (modifi ed from Neaga &

Moroz 1987).

PALAEOZOIC FROM DOBROGEA AND PRE-DOBROGEA

684

dark carbonate successions, oft en bituminous, with

thin siliciclastic interbeds. Th e lower part of this

succession (Eifelian) consists of brecciated anhydritic

rocks interbedded with shales (300 m) (Neaga &

Moroz 1987). Th e Upper Devonian is missing from

some areas, probably due to erosion. Th e Lower

Carboniferous, from the Tournaisian to the Namurian

(Serpuchovian), consists of massive limestones and

dolomites, rich in organic matter. Only the early

Late Visean is present in the eastern part of the Aluat

basin, while in the Tuzla borehole the succession is

preserved up to the base of the upper Serpuchovian

(Vdovenko 1978, 1980a, b, 1986). For the Tournaisian

a correlation with coeval deposits from the East

European Craton and Donbass Fold Belt is possible,

based on the lithology and foraminifera-dominated

microfaunal assemblages (Vdovenko 1986).

Petrographic studies (Baltres, in Roşca et

al. 1994) indicate that the carbonate rocks are

bioclastic and pelletal micrites, partly dolomitized,

with anhydrite nodules, locally associated with

sandstones and mudstones rich in brachiopod

bioclasts. Th e microfacies suggests tidal deposits,

dominated by subtidal and intertidal sediments, the

initial sediments representing bioturbated, former

limy oozes. Supratidal deposits are represented by

evaporites and suggest formation in a warm, arid

climate, in a low energy environment (lagoon or

embayment adjacent to the sea).

A very accurate biostratigraphy of the Devonian–

Carboniferous carbonate sediments east of Prut River

is based on microfauna (foraminifera, ostracods)

(Vdovenko 1972, 1978, 1980a, b, 1986; Bercenko

& Kotliar 1980) and macrofauna (brachiopods,

pelecypods, orthoceratids and bryozoans) (Safarov

& Kaptzan 1967). Th e calcareous foraminifera

belong to the Fennosarmatian province of the

North Palaeotethyan realm, characteristic of

Avalonian terranes (Kalvoda 1999). Th e main

diagnostic features of this province include: late

Frasnian diversifi ed Multiseptida-Eonodosaria-

Eogeinitzina association; late Famennian diversifi ed

Quasiendothyra association; late Tournaisian–early

Visean Kizel-Kosvin association, together with late

Visean taxa known from the southern margin of

Laurussia (Kalvoda et al. 2002).

Carboniferous Terrigenous Facies – Th e terrigenous, coal-bearing Carboniferous facies (Upper Visean–Namurian) unconformably overlies Lower–Middle Devonian sediments in the Aluat and Sărata-Tuzla basins. Th e sequence consists of grey sandstones and siltstones, with anthracite, interpreted as lacustrine sediments (Neaga & Moroz 1987). Th ey are dated as late Visean–middle Namurian based on ostracods recovered from thin limestone interbeds occurring at the top of the succession.

In the Bârlad depression, the equivalent of this coal-bearing succession is the Matca Formation. Th is is a sequence of black mudstones and grey sandstones, with local conglomerate layers. Pelitic intervals contain thin spongolithic interbeds, indicating a shallow marine environment with depths of 60–100 m. Th e rocks are rich in pyrite and coalifi ed material, suggesting reducing conditions of the basin. A proximal shelf facies developed at Matca, dominated by conglomerates and sandstones (submerged delta), and a typical distal shelf succession, dominated by clays and sandstones with sponges occurs at Burcioaia (Pană 1991). Clast petrography includes volcanic and vein quartz, plagioclase feldspars and various lithoclasts: basalts, rhyolites, spongoliths, siliceous shales, quartzites, seldom limestones; detrital micas are abundant, heavy minerals are tourmaline and opaques; framboidal pyrite repaces Endothyra tests.

Based on marine, nektobenthonic shelf fauna, with conodonts, endothyracae, ostracods, sponges, fi sh bone fragments and teeth, the dark clastic successions of the Matca Formation are assigned to the Lower Carboniferous (Middle Tournaisian–Lower–Middle Visean) (Pană 1991). Th e younger age of the Carboniferous clastics in the Bârlad depression suggests that detrital Carboniferous sedimentation has been established earlier in the western part of the Pre-Dobrogea area than in the eastern part of the Sărata basin, which might further suggest north-eastward migration of the Carboniferous basin depocentre.

Permian – Continental red-beds, associated with evaporites or volcanic-volcaniclastic rocks, represent the infi ll of the Aluat and Sărata-Tuzla basins. In the Sărata-Tuzla basin, the succession has a maximum thickness between 2000 m to over 2500 m. In the

A. SEGHEDI

685

eastern part of the Aluat graben, east of the Prut River, the Permian is dated by palaeontological and palynological data (a phyllopode association with Pseudoestheria, as well as assemblages of spores) (Kaptzan & Safarov 1965, 1966). Th e Permian age of the volcano-sedimentary successions from the Sărata-Tuzla basins is based on geochronological evidence (K-Ar ages) (Neaga & Moroz 1987), as well as on facies and geometric criteria.

Th e Permian deposits, reworking Devonian and Lower Carboniferous limestones, show abrupt facies changes along and across the basin, with lateral facies variations, from coarsest breccias and fanglomerates accumulated along the northern margin of the basin, to conglomerates and sandstones which grade southward to evaporate-bearing thin laminated siltstones and sandstones (Seghedi et al. 2003).

Th e deepest part of the Aluat basin is fi lled with thick successions of fi ne-grained continental red-beds, rich in anhydrite and gypsum. Th ey interfi nger with the coarse fanglomerates and represent deposition in playa lake and coastal sabkha environments (Seghedi et al. 2001). In Romania, their age was ascribed to the Permian, not precluding the possibility to include the Lower Triassic (Paraschiv 1986b). Below the Middle Triassic, a similar sequence of evaporitic red-beds was intersected by boreholes in the Danube Delta (Lower Prut Graben), unconformably overlying Devonian dolomitic limestones (Pătruţ et al. 1983).

In the Sărata-Tuzla grabens, the Permian succession is largely dominated by volcanic and volcaniclastic deposits, including lava fl ows and pyroclastic sediments, interbedded with continental red-beds. Based on geometric criteria, the volcanic successions are considered Lower Permian, while the evaporitic, thin laminated red-beds represent the Upper Permian, the Rotliegendes, possibly including the Lower Triassic.

Th e Permian volcanic activity yielded thick volcanic-volcaniclastic successions interbedded with continental red-beds. Volcanic products ascribed to the Lower Permian and consisting of lava fl ows and pyroclastic sediments are trapped in the Sărata-Tuzla basin, but they are known also along the north-eastern margin of the Aluat basin. Detailed facies analysis of cores recovered from borehole 1S Furmanovka revealed several superimposed upward-

fi ning cycles of alkali basalts, coarse tuff s, trachyte fl ows and ignimbritic rhyolites and rhyolitic tuff s, separated by red sandstone intervals (Seghedi et al. 2001).

Th e dominant volcanic rocks are subalkaline: subalkali basalts, hawaiites, mugearites, latites, trachytes, trachy-dacites, trachy-rhyolites (Moroz & Neaga 1996), belonging to a basalt-trachyte bimodal association (Seghedi et al. 2001). Th eir subalkaline geochemical signature (Neaga & Moroz 1987; Moroz & Neaga 1996), as well as their interfi ngering with red, continental sediments, indicate that they are products of continental, intraplate volcanism, related to extensional (possibly transtensional) rift ing. Th e Permian syn-rift volcanism was subaerial, dominantly eff usive and subordinatly explosive, as indicated by the features of the preserved volcanic products.

In contrast to the Sărata-Tuzla graben, in the eastern part of the Aluat basin intrusive rocks occur (shonkinites, alkali syenites, syenites, monzonites, granodiorite porphyries, quartz-bearing porphyritic syenites), as well as dyke rocks of the syenite group (Moroz & Neaga 1996).

In the western part of the Bolgrad-Chilia high, magnetic anomalies represent ultrapotassic magmatic bodies as confi rmed by boreholes. Such bodies vary from few metres to 20–30 m, or are most probably stocks with thicknesses over 300 m. Tectonic control of emplacement is suggested by their location at the intersection of the two main cross-cutting fault systems (Moroz & Neaga 1976; Moroz & Neaga 1996). Th e Permian age of the ultrapotassic magmatic bodies was determined from fi eld relations, as they intrude all their pre-Mesozoic country rocks, producing considerable thermal and metasomatic eff ects, especially in carbonate lithologies. Some magmatic rock samples yielded Permian K-Ar cooling ages (248 Ma) (Moroz 1984).

Palaeozoic Formations of North Dobrogea

Th e Hercynian folded basement is exposed mainly in the western, Măcin zone and forms isolated, smaller outcrop areas in the Triassic, Tulcea zone (Figure 3). Th e distribution of the Palaeozoic basement and its Cimmerian structures are shown in the map from

PALAEOZOIC FROM DOBROGEA AND PRE-DOBROGEA

686

Figure 6 and on the geological sections from Figure

7. In the Măcin zone, the Hercynian basement is

exposed in the deeply eroded cores of several NW–

SE-trending Cimmerian thrust folds. East of the

Luncaviţa-Consul Fault, the basement crops out in

the cores of two E–W-trending Cimmerian folds –

the Mahmudia anticline in the north and the Somova

anticline in the south (Murgoci 1914; O. Mirăuţă

1966b). Other smaller exposures of the pre-Triassic

basement are scattered throughout the Triassic –

Jurassic deposits of the Tulcea zone.

Two types of Palaeozoic successions were distinguished (Seghedi & Oaie 1995; Seghedi 1999) (Figure 10), joined along the Teliţa Fault, a strike-slip fault concealed by the overlying Triassic deposits. Th is fault was also inferred from geophysical evidence (Visarion & Neaga 1997) (Figures 6 & 10). Th e vertical succession of facies is shown in Figure 11.

Th e Tulcea-type Successions – Deep marine sediments, ascribed to the Ordovician–Devonian interval, have

Figure 10. Mesozoic subcrop map showing the distribution of the Palaeozoic formations in North Dobrogea, based on outcrop,

borehole and geophysical information (modifi ed from Seghedi 1999). Note that the NW–SE structural grain of the

Palaeozoic deposits and intrusions is the result of Cimmerian deformation.

A. SEGHEDI

687

Figure 11. Lithological chart for the Palaeozoic formations of North Dobrogea (modifi ed from Seghedi 1999).

PALAEOZOIC FROM DOBROGEA AND PRE-DOBROGEA

688

been separated as the Dealul Horia, Rediu and Beştepe formations (Patrulius et al. 1973, 1974) and dated using conodonts, chitinozoans and acritarcha. Devonian deposits are exposed discontinuously in the core of the Cimmerian anticline, developed south of the Sfântu Gheorghe distributary between Tulcea and Mahmudia. Th e presence of the Ordovician and Silurian deposits in the core of the Somova Cimmerian anticline, beneath the Triassic deposits, is documented in the Movila Săpată and Marca boreholes (Seghedi, in Baltres et al. 1988) (Figure 7). Th e Palaeozoic deposits form upward-coarsening sequences of radiolarian cherts to turbidites, representing northward younging, fault-bounded successions (Figure 11). Th ey are supposed to have accumulated on oceanic crust or partly on passive continental margin.

Dealul Horia Formation – A turbidite succession is exposed in Dealul Horia, south-west of Tulcea, plunging eastwards beneath the black banded cherts of the Rediu Formation. Th ese sandstone-dominated turbidites represent a succession of coarse- and fi ne-grained, greenish sandstones and siltstones. Sedimentological features indicate proximal turbidites. Sedimentary structures are partly obscured by an E–W-trending, steeply dipping penetrative slaty cleavage and phyllosilicate foliation (Seghedi & Rădan 1989). Th e deposits are ascribed to the Ordovician based on their geometric position below the Silurian Rediu Formation (O. Mirăuţă 1966b). Based on palynological assemblages dominated by acritarchs with subordinate chitinozoans, this formation was ascribed to the Upper Ordovician–Silurian (Visarion, in E. Mirăuţă et al. 1986).

Rediu Formation – Both in outcrops and boreholes, two members could be identifi ed within the Rediu Formation, corresponding to the stratigraphy established by O. Mirăuţă (1966b): a lower, siliceous member and an upper member of black or grey slates. Th e siliceous pelagic rocks are rich in silicifi ed, undeterminable radiolarians. Rocks are derived from the very low-grade metamorphism of a lithological association of black shales and radiolarian cherts (Seghedi et al. 1993). Th e age of the Rediu Formation was ascribed to the Silurian (O. Mirăuţă 1966b),

based on a rich association of conodonts, identifi ed by Elena Mirăuţă in grey limestones interbedded in the black siliceous cherts (Table 1). Th e conodont assemblage is associated with Glomospira sp., Lituotuba sp., scolecodonts, ostracods and crinoids.

Detailed palynological studies in the Movila Săpată borehole revealed an association of chitinozoans and acritarchs indicating a Lower Silurian age for the black slates (155 m thick) and a Middle–Upper Ordovician age for the underlying bedded cherts and greenish siliceous slates (350 m thick). In the Movila Săpată borehole, the associations are dominated by Chitinozoans, with scarce Acritarcha (Table 1) (Vaida & Seghedi 1996). Th e black slates from the upper part of the Rediu Formation intercepted in the Marca borehole yielded only Silurian palynological associations (Vaida & Seghedi 1999), in good agreement with the conodont-dominated microfauna previously described in outcrops. Again chitinozoans dominate the assemblage while Acritarcha are rare (Table 1).

Th e Devonian Succession – Beştepe Formation – Th e dominantly siliceous Devonian deposits are discontinuously exposed on the southern bank of the Danube in the core of the E–W-trending Triassic anticline developed between Tulcea and Mahmudia. On the northern bank of the Danube, in Ukraine, this anticline continues in the Cartal-Orlovka area, where the Devonian succession, separated as the Orlovka series, was correlated with the Beştepe Formation (Slyusar 1984).

Based on stratigraphical and palaeontological studies, the stratigraphy proposed for the Devonian deposits from the Beştepe Hills includes a lower, fl ysch-type member, a middle member made of limestones and schists and an upper member of siliceous shales (O. Mirăuţă & E. Mirăuţă 1965a, b; O. Mirăuţă 1967). A diff erent stratigraphic succession is suggested by sedimentological and structural studies: a lower member of siliceous rocks (cherts, siliceous shales with scarce, thin pelagic limestone interbeds) and an upper member of distal turbidites (Oaie & Seghedi 1994). Th e structural style of these deposits is characterized by recumbent folds and thrusts (Figure 12a) formed as result of N–S compression. A slight southward increase in metamorphic grade

A. SEGHEDI

689

Tab

le 1

.

Th

e fo

ssil

rec

ord

id

enti

fi ed

in

th

e T

ulc

ea-t

ype

Pal

aeo

zoic

dep

osi

ts f

rom

No

rth

Do

bro

gea

(aft

er E

len

a M

irău

ţă,

in M

irău

ţă 1

96

6a,

b;

Mir

ăuţă

19

71

; V

aid

a &

Seg

hed

i

19

96

; Vai

da,

in

Seg

hed

i et

al.

19

99

).

Sta

ge

Co

no

do

nts

Ch

itin

ozo

ans

Acr

itar

cha

Oth

er

D E V O N I A N

Beştepe Formation

Aco

du

s sp

.

An

gulo

du

s w

alra

thi

Bel

odel

la t

rian

gula

ris

B. d

evon

icu

s, B

. ma

crod

entu

s

Bri

anto

du

s

Hin

deo

del

la a

du

nca

, H. a

ust

inen

sis,

H. g

erm

ane,

H. p

risc

illa

Lig

onod

ina

mu

ltid

ens

L. t

hor

ia

Ois

tod

us

curv

atu

s

Oza

rkod

ina

con

gest

a

Pal

mat

odel

la d

elic

atu

lla

Pan

der

odu

s gr

aci

lis,

P. u

nic

osta

tus

P. a

ngu

stp

enn

ata

Pol

ygn

ath

us

dec

oros

a, P

. dob

roge

nis

P. e

ifl i

a, P

. fol

iata

, P. k

ocke

lian

a,

P li

ngu

ifor

mis

, P. p

enn

ata

,

Pri

onio

din

a a

lata

, P. a

lter

nat

a

Rou

nd

ya a

uri

ta

S I L U R I A N

Rediu Formation

Rediu

Formation

Oza

rkod

ina

fu

nd

amen

tata

Oz.

cf.

med

ia

Oz.

typ

ica

den

kman

ni

Neo

prio

nod

us

bicu

rvat

oid

es

Icri

odu

s sp

.

Car

nio

du

s cf

. car

nu

lus

Pal

tod

us

un

icos

tatu

s

P. c

f. r

ecu

rvat

ur

On

eoto

du

s sp

.

Aco

du

s sp

.

Con

och

itin

a e

lega

ns

EIS

.

Cya

thoc

hit

ina

fu

sifo

rmis

BO

UC

HE

C. n

ovem

pop

ula

nic

a T

AU

G.

An

goch

itin

a s

p

Lop

hos

pha

erid

ium

sp

.G

lom

ospi

ra s

p.

Lit

uot

ub

a s

p.

sco

leco

do

nts

ost

raco

ds

crin

oid

s

LA

TE

OR

DO

VIC

IAN

Lag

enoc

hit

ina

deu

nffi

P

AR

IS

Cya

thoc

hit

ina

cam

pan

ula

efor

mis

(EIS

.) A

ngo

chit

ina

lon

gico

lla

EIS

.

Mu

ltip

lici

sph

aer

idiu

m

rad

icos

um

LO

EB

LIC

H

PALAEOZOIC FROM DOBROGEA AND PRE-DOBROGEA

690

from subgreenschist facies up to lower greenschist

facies was recorded in the siliceous shales (Seghedi

1999).

Th e ‘fl ysch-like member’ is represented by tightly

folded, fi ne grained, distal turbidites, fi lling small

synclines on top of siliceous rocks (Oaie & Seghedi

1994). Th ey are made of amalgamated thin beds

showing Tcde and Tde Bouma divisions (horizontal

and ripple cross-laminated, fi ne-grained calcareous

sandstones and dark mudstones) (Figure 12b), with

abundant fl ute casts at the lower part of the sandstone

lithofacies (O. Mirăuţă 1967). Ichnofauna is abundant

at the interface between d and e Bouma divisions and

develops on top of the pelitic lithofacies (Figure 12c).

Th e ichnofaunal association with Helminthoides,

Chondrites, Protopalaeodyction, etc., belongs to the

Nereites ichnofacies and suggests accumulation at

the distal parts of the turbiditic fans, at water depths

exceeding 800 m (Oaie 1989). Sedimentological

features suggest that the turbidites accumulated in a

sediment-starved, ponded basin.

Th e siliceous member represents pelagic deposits,

derived from radiolarian cherts and muds (Seghedi

et al. 1993). Th ey are interpreted to have accumulated

on oceanic basement (Seghedi & Oaie 1994).

Petrographic studies revealed that the lithological

association is dominated by bedded radiolarian chert

(Figure 13) and siliceous slates, with minor bedded

Figure 12. Views from the Tulcea-type Palaeozoic. (a) Outcrop-scale low-angle thrust in bedded cherts of Beştepe Formation (eastern

part of the Beştepe Hills). (b) Tight normal folds in distal turbidites of Bestepe Formation (Pârlita abandoned quarry,

Victoria). (c) Meandering trails on top of the pelitic lithofacies from distal turbidites of Beştepe Formation (Ilgani, Nufăru

village). Th e trails indicate the deep water Nereites ichofacies.

(a)

(b)

(c)

A. SEGHEDI

691

iron carbonates and bedded iron sulphates and black slates (Seghedi et al. 1993). Th ey show mineralogical evidence for deposition in anoxic conditions in a basin below the CCD. Local interbeds of pelagic limestones suggest that for some time intervals, deposition took place above the CCD level. Pelagic limestones form thin (1–5 cm) or thicker (10–40 cm) layers, interbedded locally in the siliceous rocks. Th ey have not been intercepted in any of the deep boreholes drilled in the core of the Tulcea-Mahmudia anticline.

Th e Devonian age of the Beştepe Formation is based on the conodont fauna identifi ed in the interbedded limestones from the Beştepe Hills and a

few other outcrops (O. Mirăuţă 1967, 1971). Similar micritic limestones, rich in more or less recrystallized microfauna and upper Devonian conodonts (Asejeva et al. 1981), are interbedded with siliceous deposits from the Orlovka quarry, on the northern bank of the Danube (Slyusar 1984). Clastic deposits at Orlovka yielded a Devonian palynological association (Velikanov et al. 1979), in good agreement with the conodont ages.

Th e Carboniferous is not documented in the deep marine sediments, although an indication for the Lower Carboniferous was given by palynological data mentioned above. Considering the structure of these successions, with northward younging tectonic

Figure 13. Views from the cherts of the Beştepe Formation. (a) Borehole core of bedded chert (ch), consisting of white and dark

siliceous layers rich in radiolaria (r). (b) Chert layer with radiolaria replaced by secondary silica. (c) Chert with ghosts of

silicifi ed radiolaria, one individual still preserving spikes.

(a)

(b)

(c)

PALAEOZOIC FROM DOBROGEA AND PRE-DOBROGEA

692

units formed by tectonic accretion beneath the upper

plate, the presence of the Carboniferous cannot be

precluded at depth, as shown in Figure 7.

Th e Măcin-type Successions

Th e largest area of Palaeozoic basement in North

Dobrogea includes pre-Silurian metamorphic

formations, Silurian–Upper Palaeozoic

anchimetamorphic rocks, as well as calc-alkaline

volcaniclastic suites and granitoids (Figure 11).

Th e relationships between distinct metamorphic

terranes are tectonic, as well as their relations to the

fossil-bearing Palaeozoic deposits (Seghedi 1980,

1986a). Th e Silurian to Upper Palaeozoic formations

were deformed in very low-grade metamorphic

conditions, developing a slaty cleavage, penetrative in

most lithologies (Seghedi 1985, 1986b). In Silurian–

Lower Devonian lithologies, the steeply-dipping

penetrative cleavages strongly obliterate bedding.

Th is deformation occurred prior to emplacement of

granitoid intrusions, which are surrounded by large

areas of contact metamorphism.

Metamorphic formations belong to three thrust-

bounded groups, which vary in metamorphic grade

from middle-upper amphibolite facies (Orliga

and Megina terranes) to lower greenschist facies

(Boclugea terrane) (Figure 7). Th e lithology of the

Orliga terrane is dominated by micaceous quartzites,

with subordinate metapelites, metabasic rocks and

crystalline limestones, interpreted as an ancient

accretionary complex (Seghedi & Oaie 1995; Seghedi

1999). Metabasites in the Orliga Groups have the

geochemical characteristics of ocean fl oor tholeiites

(Crowley et al. 2000). Quartzites and phyllites make

up the Boclugea Group, a low-grade series of biotite

grade metasediments. Th e Megina Group, dominated

by amphibolites, with minor acid metavolcanics

and metapelites, is associated with orthogneisses.

Geochemical features indicate that amphibolites

represent ocean fl oor tholeiites associated with calc-

alkaline rhyolitic volcanics (Seghedi 1999). REE

geochemistry suggests that mafi c volcanics of the

Orliga and Megina terranes were generated by partial

melting of a variably depleted mantle asthenosphere,

with a contribution from a continental lithosphere

component (Crowley et al. 2000).

Based on fi eld relations, metamorphic grade and K-Ar geochronology, the ages of the metamorphic suites have been variously ascribed to the Devonian (Murgoci 1914; Rotman 1917), Cambrian and Ordovician (O. Mirăuţă 1966a; Seghedi 1980), or Late Precambrian (Ianovici & Giuşcă 1961; Giuşcă et al. 1967; Kräutner et al. 1988; Seghedi 1980, 1986a). An Early Cambrian age was assigned to the muscovite schists of the Boclugea group based on palynological data (Vaida & Seghedi 1999). Th e microfl oral association, yielded by samples from the Iulia borehole, consist of Acritacha, including Micrhystridium sp., M. lanatum, Leiomarginata simplex, Granomarginata prima, Uniporata nidius, Tasmanites sp.

Th e youngest detrital zircon U-Pb ages (500–536 Ma) yielded by metamorphic rocks from the Măcin zone indicate that the maximum depositional age for the sedimentary protoliths of both the Boclugea and Orliga sediments is Middle to Late Cambrian (Balintoni et al. 2010).

Th ere are several lines of evidence indicating the age of the Hercynian metamorphism. Single grain Ar-Ar dating of muscovites from Orliga samples has yielded early Permian ages of 275±4 Ma on micaschist and 273±4 Ma on paragneiss (Seghedi et al. 1999) (Figure 8). Monazite CHIME dating of metapelites from the Orliga terrane yielded ages ranging between 326±25 Ma and 282±38 Ma, indicating that the amphibolite facies metamorphic event took place in the Late Carboniferous–Lower Permian time span (Seghedi et al. 2003a). Similar age ranges were yielded by monazites for the Megina metapelites (Seghedi, Spear & Storm Hitchkock, unpublished data).

Th e Silurian – Cerna Formation – Th e succession includes a lower member of dark grey limestones and shales, rich in pyrite and formed in an euxinic environment, and un upper member of black argillites (O. Mirăuţă & E. Mirăuţă 1962; O. Mirăuţă 1966a); the argillites show interbeds of quartz-muscovite sandstones and brown-yellowish limestones with rare organic debris (unidentifi ed gastropods, crinoid ossicles). In their main outcrop area, the argillites show highly discontinuous, lens-shaped bodies of magmatic rocks with centimetric-decimetric thicknesses. Because these rocks are aff ected by strong hydrothermal alteration, identifi cation

A. SEGHEDI

693

of their protoliths is extremely diffi cult. In most outcrop areas, the Silurian deposits are overthrust by quartzites and phyllites of the Boclugea group (O. Mirăuţă 1966a; Seghedi 1986a). Th e tight folding of the deposits and superimposed deformation makes thickness estimation diffi cult. From bottom to top, the facies succession indicates that basin shallowing was accompanied by a transition from restricted to normal marine conditions.

Th e Cerna Formation was loosely ascribed to the Silurian based on macrofaunal remains, like Cyathophyllum corals (Simionescu 1924), a Rastrites fragment (O. Mirăuţă & E. Mirăuţă 1962), the conodont Panderodus sp. and scarce debris of tentaculites and corals (Iordan 1999). Th e identifi cation of the Lower Devonian conodont Icriodus woschmidti (O. Mirăuţă 1966a) in brown limestone interbeds within the argillitic succession suggests an initial lithological continuity across the Silurian/Devonian boundary.

Th e Lower Devonian – Bujoare Formation – Th e Lower Devonian deposits of the Măcin zone show a thickness of 300–400 m (O. Mirăuţă & E. Mirăuţă 1962) and develop in Rhenish facies (Iordan 1974). Th e main lithological types are grouped into two members, representing the Geddinian and Coblenzian (O. Mirăuţă 1966a), respectively the Lochkovian–Emsian. According to the stratigraphy of these authors, the lower part of the Lower Devonian succession includes white and grey limestones, overlying quartzitic sandstones (Figure 14a) interbedded with dark slates and brown crinoidal limestones with Icriodus woschmidti. Th e upper part of the succession consists of discontinuous beds of quartzitic sandstones, black slates, calcareous limestones and crinoidal limestones (O. Mirăuţă & E. Mirăuţă 1962). Th e fauna frequently forms coquina beds, suggesting deposition in an open shelf benthic environment (Iordan 1999).

Th e age of the clastics is indicated by a rich brachiopod-dominated fauna recovered from Bujorul Bulgăresc Hill (Cădere & Simionescu 1907; Simionescu 1924), attesting the presence of the Pragian–Emsian (Iordan 1974). Besides brachiopods (Figure 14b), the Lower Devonian fauna contains crinoids and tentaculitids, with subordinate trilobites, corals, bryozoans and ostracods (Iordan 1974).

A review of the Lower Devonian fauna (Iordan 1999) revealed several features of the main faunal assemblages presented in Table 2. Th e brachiopods form coquinas, along with corals, bryozoans, ostracods and trilobite fragments (Table 2). Tentaculitids also form coquinas, sometimes exclusively covering the rock surfaces and frequently associated with crinoids (Iordan 1974).

Th e Late Palaeozoic – Carapelit Formation – Continental deposits, reworking granites, quartzites and phyllites (Murgoci 1914; Rotman 1917), were separated as the Carapelit formation (Mrazec & Pascu 1896). Th eir primary relations with the metamorphic basement are obscured due to subsequent Cimmerian deformation. Th e oldest exposed conglomerates rework limestone clasts that yielded Middle–Upper Devonian conodonts (O. Mirăuţă & E. Mirăuţă 1962), while younger conglomerates rework quartzite and granite clasts, suggesting a reverse clast stratigraphy. Th e stratigraphic succession of the Carapelit Formation (Figure 11) consists of lower, grey alluvial deposits, followed by continental red-beds and an upper volcano-sedimentary succession (Oaie 1986; Seghedi & Oaie 1986; Seghedi et al. 1987).

Alluvial deposits consist of grey alluvial fan-alluvial plain sequences, with debris fl ow and stream fl ood conglomerates dominating the coarse members and sandstone-siltstone cycles in the fl ood-plain deposits (Seghedi & Oaie 1986) (Figure 14c).

Red beds (Martina red sandstones), up to 900 m thick, comprise a succession of pebbly red sandstones (Figure 15a), showing both horizontal and planar cross stratifi cation, and organized conglomerate beds, with locally preserved clast imbrication; dessication cracks and raindrop imprints are preserved in places in thin, clayey, purple mud-drape facies (Oaie 1986). Th ere is good fi eld evidence that the red beds directly overlie wedge-shaped, alluvial fanglomerates (Seghedi et al. 1987). Vertical facies distribution indicates an upward-coarsening sequence, deposited by a sandy braided river with fl uctuating discharge, showing upward progradation of coarse, longitudinal bar deposits over sand dunes (Oaie 1986). Sandstone petrography suggests that the onset of red bed deposition was related to a major climatic change, switching from a warm and humid climate which

PALAEOZOIC FROM DOBROGEA AND PRE-DOBROGEA

694

prevailed during alluvial fan sedimentation, to

an arid, dry climate, that controlled the red bed

accumulation (Seghedi & Oaie 1986; Oaie 1986).

Th e upper part of the Carapelit Formation

is dominated by thick volcano-sedimentary

successions, consisting of superimposed cycles

of pyroclastic deposits and coarse rhyolitic

epiclastic sequences. Pyroclastic rocks (Figure

15b) are dominated by large volumes of calc-

alkaline ignimbritic rhyolites (up to 1000 m thick),

interbedded with air fall tuff s, rare base surge

deposits or accretionary lapilli tuff s and display the

geometry of superimposed fl ow units. Vertical facies

associations suggest that the style of sedimentation

was controlled by intermittent volcanic eruptions

(Seghedi et al. 1987). Overall volcanological

features indicate that Hercynian volcanism in North

Dobrogea was subaerial, characterized by calderas

and plinian eruptions, while accumulation of the

volcanic products was both subaerial and subaqueous

(in fl uvial and lacustrine environments).

Due to lack of fossils, the age of the Carapelit

Formation is poorly constrained. It has been ascribed

to the Permo–Carboniferous (Mrazec & Pascu 1986;

Rotman 1917), Carboniferous (Cantuniari 1913),

Lower Carboniferous–Upper Devonian (Murgoci

(a)

(c)

(b)

Figure 14. Views from the Macin-type Palaeozoic. (a) Concentric folds deform decimetric orthoquartzite bed in the lower Devonian

Bujoare Formation. Note cleavages fanning in anticline hinge. Only penetrative, steeply dipping cleavage planes are seen

in mica-rich fi ne-grained sandstones overlying the orthoquartzite bed (Chior Tepe). (b) Brachiopod coquina in Bujoare

Formation (Bujorul Bulgăresc Hill; aft er Iordan 1974). (c) Steeply-dipping slaty cleavage planes obliterating bedding in the

alluvial plain member of the Carapelit Formation (Amzalar Hill).

A. SEGHEDI

695

Tab

le 2

. Th

e

pal

eon

tolo

gica

l re

cord

of

the

Măc

in-t

ype

Pal

aeo

zoic

fo

rmat

ion

s (a

ft er

Io

rdan

19

74

; Mir

ăuţă

19

65

; Vai

da,

in

Bal

tres

et

al.

19

88

)

Lo

wer

Dev

on

ian

Bu

joar

e

Fo

rmat

ion

Bra

chio

po

ds

Sch

izop

hor

ia p

rovu

lvar

ia (

Mau

r).

Dal

man

ella

cir

cula

ris

(So

w),

D. f

asc

icu

lari

s (d

’Orb

.)

‘Ort

his

’ str

igos

a S

ow

.

Lep

taen

a s

p. e

x gr

. rh

om

bo

idal

is (

Wil

ch.)

Ph

olid

ostr

oph

ia a

ff . n

aran

joan

a (

Ver

n.)

Rh

enos

trop

hia

su

bar

ach

noi

dea

(A

rch

.et

Ver

n.)

Stro

pheo

don

ta a

ff . s

edgw

icki

(A

rch

.et

Ver

n.)

S. e

xpla

nat

a (

So

w.)

Dou

vill

ina

in

ters

tria

lis

(Ph

ill.)

Sch

ellw

ien

ella

hip

pon

ix (

Sch

nu

r)

Ch

onet

es s

arci

nu

latu

s (S

chl.)

C. p

leb

eju

s S

chn

ur

Stro

phoc

hon

etes

ten

uic

osta

tus

(Oeh

lert

)

Ret

ich

onet

es a

ff . m

inu

tus

(Bu

ch)

Acr

ospi

rife

r pr

ima

evu

s (S

tein

.)

A. p

elli

co (

Gei

b.)

A. p

elli

co p

rim

aev

ifor

mis

(S

cup

.)

Eu

rysp

irof

er a

rdu

enn

ensi

s (S

chn

ur)

Hys

tero

lite

s h

yste

ricu

s (S

chl.)

Bra

chys

piri

fer

cari

nat

us

(Sch

nu

r)

Fim

bris

piri

fer

bisc

hofi

(R

oem

.)

‘Spi

rife

r’ c

f. u

nd

uli

fer

Kay

s.

Un

cin

ulu

s cf

. an

tiqu

us

(Sch

nu

r)

Cam

arot

oech

ia s

p.

Meg

ante

ris

cf. a

rch

iaci

(V

ern

.)

Co

rals

Cya

thop

hyl

lum

ex

gr. C

.

rad

icu

lum

Ro

m.

Lin

dst

roem

ia c

f. d

alm

ania

Pae

ck.

Zap

hre

nti

s sp

.

Cla

doc

hon

us

sp.

Ten

tacu

liti

ds

Ten

tacu

lite

s sc

alar

is

Sch

l.

Un

icon

us

du

rus

Lu

dw

.

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no

ids

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us

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hu

s sp

.

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odu

s w

osch

mid

ti

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ozo

ans

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dot

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a s

p.

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este

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sp

.

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itry

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sp

.

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lob

ites

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leti

na

cf.

ham

mer

sch

mid

ti (

Ric

ht)

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alo

not

us

sp.

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aco

ps s

p.

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uri

anC

ern

a

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rmat

ion

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thop

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no

do

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der

odu

s sp

.

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pto

lite

s R

ast

rite

su

nid

enti

fi ed

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tro

po

d

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ovi

cian

Cam

bri

an

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clu

gea

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artz

ites

(Iu

lia

bo

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ole

)

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yno

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s

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rhys

trid

ium

sp

.

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anat

um

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omar

gin

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.

PALAEOZOIC FROM DOBROGEA AND PRE-DOBROGEA

696

1914) and Lower Carboniferous (O. Mirăuţă & E. Mirăuţă 1962). Th e only fossil discovered so far is a plant, variously identifi ed by various authors as Calamaria, Asterocalamites or Archaeocalamites (Atanasiu 1940); these taxa indicate various time intervals – Upper Carboniferous or Upper Devonian–Lower Carboniferous. A re-examination of this fossil is impossible, as the sample seems to be lost. Pelitic samples from Carapelit Hill yielded a poorly-preserved palaeontological content, obviously reworked, including acritarchs, trilete spores, chitinozoans, scolecodonts and microspores, all suggesting the Silurian–Carboniferous interval (Vaida, in Seghedi et al. 1986). All existing information, derived from the petrography of coarse-

grained members, palynological data and microfaunal

content of the reworked limestone clasts indicate that

the Carapelit Formation is certainly post-Devonian.

A Permo–Carboniferous (Late Carboniferous–early

Permian) age is favoured, based on the presence of

red beds and volcaniclastic sedimentation (Mrazec &

Pascu 1986; Rotman 1917; Oaie & Seghedi 1992).

Hercynian Calc-alkaline Intrusives – Field relations

demonstrate that two main phases of granitoid

emplacement took place in the Măcin Zone, before

and aft er the deposition of the Carapelit Formation

(Rotman 1917). Older generation intrusives include

mainly biotite granite, porphyritic biotite granite or

Figure 15. Views from the Carapelit Formation and Permian dykes. (a) Steep bedding in red sandstone and microconglomerate of

the Carapelit Formation (Crapcea Hill). (b) Rhyolitic ignimbrite of the Carapelit Formation, showing pumice (dark) and

crystalloclasts (light) in a welded tuff matrix. (c) Steep dykes of trachyte (left ) and alkali basalt (right) emplaced in turbidites

of the Bestepe Formation (Pârlita abandoned quarry, Victoria). (d) Detail from the trachyte dyke, showing K feldspar

phenocrysts in a pink, fi ne-grained matrix (Pârlita abandoned quarry, Victoria).

(a)

(c)

(b)

(d)

A. SEGHEDI

697

leucogranite, emplaced in Boclugea quartzites and

phyllites and accompanied by contact metamorphic

aureoles or metasomatic eff ects resulting in garnet-

diopside skarns. Such granites, along with their low-

grade metamorphic host rocks are reworked in the

Carapelit conglomerates.

Younger generation intrusives represented by

the Greci massif form a suite of highly diff erentiated

I-type calc-alkaline rocks, ranging from dioritic and

gabbroic facies to leucogranites, but dominated by

biotite-hornblende granodiorites and tonalites. Th ese massifs are emplaced in the Carapelit Formation as high level, high temperature intrusives, with both cross-cutting and partly conformable contacts, producing large contact metamorphic areas and local garnet-pyroxene skarns. Rocks are rich in xenoliths of hornfelsed country rocks, as well as various types of cognate xenoliths. Geochemical data suggest a mantle source for basic end members and a deep or mid-crustal environment for the genesis of the granitic rocks, as well as evolution of a syn-collisional geotectonic setting (Tatu & Seghedi 1999).

K-Ar geochronological ages indicate Late Carboniferous (306, 320, 329 Ma) and early Permian ages (264 Ma) for the main intrusive bodies (Mînzatu et al. 1975). An apparent plateau age of 242.2±2.1 Ma, yielded by biotite concentrate from the Greci granodiorite, indicates the Late Permian-Skythian interval (Seghedi et al. 1999). Considering that these intrusives are cut by early Triassic dolerite and rhyolite dykes, this thermal event has been interpreted as the result of ductile mylonitization of the Hercynian basement during early Triassic rift ing.

Intrusives in the Tulcea Zone, dated as Middle Carboniferous by the K-Ar method (Mînzatu et al. 1975) represent a suite of peraluminous, S-type diff erentiated granites, emplaced as high temperature, low level intrusions (Seghedi et al. 1994b). Th eir emplacement produced a high temperature-low

pressure metamorphism in their Lower Palaeozoic

host rocks. Th is suite consists of two mica granites,

with hornblende in minor tonalitic facies and

garnet in pegmatites, commonly rich in biotitic

xenoliths and in zoned K-feldspar megacrysts. Th eir

geochemistry suggests formation by plagiclase and

zircon fractionation from a magma formed by partial

melting of the lower crust.

Late Permian Alkaline-Sub-alkaline Magmatism

Dyke Swarms of the Basalt-trachyte Bimodal Association – Alkaline magmatic rocks are recorded in two diff erent areas of North Dobrogea, forming distinct petrological associations. A tectonic control is suggested by their development south of the Sfantu Gheorghe Fault and north of the Peceneaga-Camena Fault.

An E–W-trending dyke-swarm of the basalt-trachyte bimodal association is emplaced exclusively in the Upper Ordovician–Devonian deep marine successions developed between the Sfântu Gheorghe and Teliţa Faults (Figure 9). Th e E–W- to ESE–WNW-trending, steeply-dipping dykes are known both from outcrop (Figure 15c), and from borehole drillings (Seghedi, in Baltres et al. 1988, 1991, 1992). Th e age of the dyke-swarms is pre-Triassic, as documented by fi eld relations in outcrops at Hora Tepe (Monument Hill), Tulcea (O. Mirăuţă 1966b). On the northern slope of Hora Tepe in Tulcea, steeply dipping trachyte dykes intruding Devonian siliceous sediments of the Beştepe Formation are truncated by Lower Triassic (Werfenian) fanglomerates which rework clasts of the dyke rocks.

Both in outcrops and boreholes, the subvolcanic rocks display intense hydrothermal alteration that obscures their initial petrography (Seghedi et al. 1994). In basalts and olivine basalts, olivine and pyroxenes are replaced by dolomite and iron carbonates, associated with various amounts of chlorite and secondary silica. Trachytes are porphyritic, with alkali feldspar phenocrysts in a feldspar-dominated matrix (Figure 15d) showing weaker hydrothermal alteration, with only incipient replacement of feldspars by carbonates and illite.

Geochemical features of the dykes reveal the bimodal character of the magmatism and suggest an intraplate setting (Seghedi et al. 1994a, b). Th e bimodal association and the presence of high-K trachytes suggest magma generation in a complex extensional setting, probably transitional between an active continental margin and extensional (transtensional) intraplate setting. Th e structural trend of the dyke swarm suggests generation in a N–S-oriented stress fi eld. Th e bimodal association shows similar petrographic features and geochemical signature and even the same type of extensive

PALAEOZOIC FROM DOBROGEA AND PRE-DOBROGEA

698

hydrothermal alteration as the Permian volcanic rocks from the Scythian Platform, and therefore the dykes are interpreted as connected with the Permian rift ing of the Scythian Platform.

Late Permian Sub-alkaline Magmatism – Along the south-western margin of North Dobrogea, parallel to the Peceneaga-Camena Fault, several small bodies of alkaline rocks have been emplaced in the Palaeozoic country rocks, including the Carapelit Formation (Cantuniari 1913; Întorsureanu et al. 1989) (Figures 6 & 7a). Th e most representative, Turcoaia massif, is made up of two distinct plutonic and hypabyssal associations which were intruded in a narrow time range as a concentric or ring-complex (Tatu & Teleman 1997; Tatu 1999). From the core to the periphery, quartz-syenites, monzogranites, alkali-granites and alkali-rhyolites occur. Values of the 87Sr/ 86Sr ratio (Pop et al. 1985) suggest that crustal anatexis took place in a continental, intraplate setting (Tatu 1999). Mineralogical and geochemical data indicate that the Turcoaia alkaline massif represents an A-type magmatic type with complex within-plate evolution.

Th e question of the emplacement age is not yet properly answered, as whole rock Rb-Sr radiometric dating has yielded a large age range from Permian to Triassic (Mînzatu et al. 1975). However, a Late Permian Ar-Ar age was yielded by riebeckites from the Turcoaia massif (Teleman, unpublished data). As the plutonic complex is cut by lower Triassic dykes of the basalt-rhyolite bimodal association, the Ar-Ar age should be very close to the emplacement age of the Turcoaia massif.

Palaeozoic Deposits of East Moesia

Th e Palaeozoic formations of East Moesia are documented in 16 boreholes in South Dobrogea and in 19 boreholes in the Romanian Plain (Figure 16). Aft er penetrating thin Palaeozoic successions, several boreholes in the Romanian Plain bottomed in Ediacaran basement (Bordei Verde, Ţăndărei, Călăraşi, Zăvoaia boreholes etc.), which shows similar facies with that from Central Dobrogea. Based on these data, the spatial distribution of the East Moesian Palaeozoic deposits and of their basement rocks are shown in Figure 17.

Th e overall lithological succession and main biostratigraphic zones of the East Moesian Palaeozoic are shown in the synthetic lithological log from Figure 18. Detailed lithological logs of the main boreholes are presented in the review of the Moesian Palaeozoic (Seghedi et al. 2005a, b).

Th e Palaeozoic succession begins with moderately dipping Cambrian–Ordovician clastics and/or Wenlock to Lower Ludlow shales. It is considered that the last folded member of this unit is the Lower Ludlow, because in the 5083 Mangalia and 2881 Călăraşi boreholes, the Lower Ludlow is overlain above an angular unconformity by a fl at-lying succession starting with the Upper Ludlow (Răileanu et al. 1966, 1967; Paraschiv 1974).

Th e Silurian deposits, with a maximum thickness of 700 m at Călăraşi, overstep the Ordovician strata, directly overlying the Upper Neoproterozoic basement in the Romanian Plain. Th e largest part of the Llandovery corresponds to a regional gap, only its uppermost part being preserved in some boreholes (Iordan 1981, 1982, 1984). An angular unconformity was recorded between the Lower and Upper Ludlow in the Călăraşi and Zăvoaia boreholes, accompanied by a lithological change from graptolite shales to argillites (Răileanu et al. 1967; Iordan 1981).

Th e Silurian–Devonian boundary is continuous in the Zăvoaia and 2881 Călăraşi boreholes, marked by a change in facies from argillites to quartzitic sandstones. An angular unconformity was recorded between the Silurian and Devonian in the 5083 Mangalia borehole (Răileanu et al. 1966), similar to that recorded on the northern and western platform margin, in West Moesia (Paraschiv 1974).

In East Moesia the Tournaisian–Lower Visean corresponds to a regional stratigraphic gap (Iordan 1999), unlike in West Moesia, where the Upper Devonian carbonate succession continues into the Carboniferous as far as the Namurian (Lower Bashkirian) (Iordan 1988).

Cambrian–Ordovician – In South Dobrogea the presence of the Cambrian is not yet documented by fossils. It is assumed that the Cambrian represents the lower part of the shallow marine quartzitic sandstones over 500 m thick, intercepted in the 5083

A. SEGHEDI

699

Mangalia borehole, which in their upper part contain

Ordovician palynological assemblages (Paraschiv et

al. 1983).

Th e Ordovician succession consists of sandstones

and orthoquartzites with shale interbeds, overlain

by glauconite-bearing shales (Murgeanu & Patrulius

1963; Beju 1972). Th e shales contain graptolites,

palynomorphs and brachiopods. Th e Ordovician is

characterized by the hirundo graptolite zone (Iordan

1992) and by O1 and O

3 palynomorph zones (Beju

1972, 1973). In the Romanian Plain, an Arenig fauna

was recorded in black argillites and grey-greenish

glauconitic siltstones from the Bordei Verde borehole.

Th e fauna includes the graptolites Didymograptus

hirundo (Salt.) and the inarticulate brachiopods

Lingulella aff . davisii McCoy (Murgeanu & Spassov

1968).

Figure 16. Distribution of main boreholes intercepting Palaeozoic deposits in East Moesia (modifi ed from Seghedi et

al. 2005).

PALAEOZOIC FROM DOBROGEA AND PRE-DOBROGEA

700

Silurian – Th e Silurian sediments show the typical

graptolite shale facies, classical in the lower part

(Wenlock–Lower Ludlow) and mixed in the upper

part (Upper Ludlow and Přídolí) (Grigoraş 1956;

Iordan 1981, 1990; Paraschiv & Muţiu 1976). Th e

maximum total thickness of the Silurian deposits

from East Moesia is about 1000 m. In the Călăraşi

and Zăvoaia boreholes, the limit between the Lower

Figure 17. Permian subcrop map showing the distribution of Precambrian basement rocks and Palaeozoic deposits in

East Moesia (modifi ed from Seghedi et al. 2005).

A. SEGHEDI

701

Figure 18. Schematic lithological log of Palaeozoic formations from East Moesia (modifi ed from Seghedi et

al. 2005a), showing the graptolites, tentaculites, trilobites, conodonts and brachiopod zones (aft er

Răileanu et al. 1967; Iordan 1967, 1972, 1981, 1985, 1987b, 1988, 1992; Iordan & Rickards 1971; Iordan

et al. 1985, 1987a) and chitinozoan zones (Beju 1972; Vaida et al. 2004).

PALAEOZOIC FROM DOBROGEA AND PRE-DOBROGEA

702

and Upper Ludlow is marked by a change in facies from black shales to argillites. In the mixed type of the graptolite shale facies, the fauna is represented by small bivalves (Cardiolidae, Lunulacardiidae and Antipleuridae families), fl attened orthocone cephalopods, microbrachiopods, corals (Favosites gothlandicus), crinoidal ossicles and scarce graptolites (Table 3).

Th e biostratigraphy of the Silurian is based on graptolite zones: insectus-centrifugus, murchisoni and fi rmus zones for the lower Wenlock (Ţăndărei and Bordei Verde boreholes); radians and lundgreni for the upper Wenlock (Iordan & Rickards 1971); nilssoni-scanicus and incipiens for the lower Ludlow (Tuzla-Costineşti, 5083 Mangalia, Biruinţa, 2881 Călăraşi and Ţăndărei boreholes); bohemicus and inexpectatus zones for the Upper Ludlow (Biruinţa, Ţăndărei, Călăraşi and Zăvoaia boreholes) (Iordan 1981, 1990); ultimus-formosus for the Přídolí (Ţăndărei, 2881Călăraşi and Zăvoaia boreholes) (Table 4). Th e Silurian chitinozoan associations are described in the Ţăndărei borehole (Beju 1972; Iliescu 1976) and Zăvoaia (Vaida et al. 2004; Vaida & Verniers 2005), the latter including the very short range top Přídolí Eisenachitina fi lifera and E. lagenomorpha.

Devonian – Th e lower Devonian (Lochkovian –Emsian), up to 460 m thick, represents a continuation of the pelitic facies of the Silurian. It consists of argillites, argillaceous shales, marly or sandy argillites, dark siltstones with subordinate grey quartzitic sandstones, showing coquina interbeds. Th e lower Devonian–Eifelian interval (Smirna quartzitic sandstones), 120 m thick, is dominated by quartzitic sandstones, orthoquartzites and quartzitic conglomerates, with minor interbeds of black argillites and dolomitic limestones. Red arkosic sandstones occur toward the top of the sandstone facies in the 5083 Mangalia and 2881 Călăraşi boreholes.

Th e overlying Givetian–Famennian deposits represent a carbonate-evaporite succession up to 905 m thick in the Smirna borehole, separated as the Călăraşi Formation s. str. (Răileanu et al. 1967; Iordan 1981). Th e bio- and lithofacies of the Călăraşi Formation include organogenic limestones, dolomites, as well as black limestones with evaporite interbeds, with minor black argillites and sandstones.

Th e lithology refl ects various tidal-fl at environments, from supratidal to intertidal, as revealed by microfacies studies (Vinogradov & Popescu 1984).

Th e biostratigraphy of the Devonian was established in two boreholes in East Moesia, the 5082 Mangalia and 2881 Călăraşi boreholes (Răileanu et al. 1967). Besides conodonts, tentaculites and trilobites, the Devonian biostratigraphy is based on brachiopods, psilophytal plants, as well as placoderm and ostracoderm fi shes.

Th e Lochkovian is indicated by the Icriodus woschmidti conodont zone (Răileanu et al. 1967). Two global chitinozoan biozones of the Lochkovian were identifi ed: Eisenackitina bohemica and Urochitina simplex (Beju 1972; Vaida et al. 2004). Th e Lochkovian assemblage also contains Bursachitina oviformis Eisenack (Lakova 1993; Vaida et al. 2005).

Th e Pragian is indicated by the tentaculites zones Nowakia acuaria and Voliynites velainii (Iordan 1967, 1981; Iordan et al. 1985, 1987b). Besides tentaculites, the assemblage includes brachiopods, bivalves, crinoids, ostracods, corals and gastropods (Table 4).

Th e Emsian faunal assemblage is dominated by trilobites and bivalves, with subordinate brachiopods, gastropods, orthoceratidae, corals, crinoids, bryozoans and ostracods. Th ree Emsian trilobite zones were identifi ed in the Mangalia and Călăraşi boreholes: Pseudocryphaeus prostellans, Pilletina asiatica and Dipleura fornix (Iordan 1967, 1981, 1988). Th e trilobite assemblage is presented in Table 4.

Th e brachiopods recovered indicate a Lochkovian–Emsian age (Iordan 1999). Th e most representative species, along with associated bivalves, are shown in Table 4.

Th e Eifelian is characterized by the N. maureri tentaculites zone. Th e Eifelian clastic facies is rich in psilophytal plants, placoderm and ostracoderm fi shes, while brachiopods, tentaculites, crinoids, bivalves, gastropods, corals, eurypterids, ostracods and conodonts are subordinate. Psilophital plants include Pseudosporochonus krejcii Pot. et Bern., Aneurophyton germanicum Kr. et Weyl., Calamophyton primaevum Kr. et Weyl. and Hyenia sp. (Răileanu et al. 1966; Iordan et al. 1985). Th e placoderm and ostracoderm fi shes (Lunaspis broilii Gross, L. sp., ?Wijdeaspis sp.

A. SEGHEDI

703

Tab

le 3

. Th

e m

ain

fau

nas

rec

ord

ed i

n t

he

grap

toli

te s

hal

es a

nd

mix

ed f

acie

s o

f th

e Si

luri

an i

n E

ast

Mo

esia

(Io

rdan

19

81

, 19

90

), r

evis

ed b

y Io

rdan

(1

99

9).

Gra

pto

lite

sB

ival

ves

Cep

hal

op

od

s –

Ort

ho

con

e

Nau

tilo

ids

Bra

chio

po

ds/

Cri

no

ids

Pří

do

lí

Bo

reh

ole

s C

ălăr

aşi,

Zăv

oai

a an

d Ţ

ănd

ărei

Mon

ogra

ptu

s ex

gr

form

osu

s B

ou

č.

M. s

p. A

Saet

ogra

ptu

s ra

rus

(Tel

ler)

Pri

stio

grap

tus

grig

ora

şii

Iord

an

Lin

ogra

ptu

s p

osth

um

us

(Ric

hte

r)

L. s

p. 1

Car

dio

lita

cf.

boh

emic

a (

Bar

r.)

C. c

f. f

orti

s (B

arr.

)

‘Car

dio

la’ i

nso

lita

Bar

r. L

un

ula

card

ium

un

du

latu

m B

arr.

Du

alin

a s

p.

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hoc

era

s’ v

erte

brat

um

(B

arr.

)

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son

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as

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le (

Bar

r.)

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icu

loid

aea

sp

. Lep

tost

roph

ia s

p.

Ple

ctod

onta

sp

.

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elle

lla

sp

.

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talo

crin

us

sp. S

cyph

ocri

nit

es s

p.

Up

per

Lu

dlo

w

Bo

reh

ole

s C

ălăr

aşi,

Z

ăvo

aia,

Ţăn

dăr

ei a

nd

Bir

uin

ţa

Boh

emog

rapt

us

boh

emic

us

(Bar

r.)

Pri

stio

grap

tus

grig

ora

şii

Iord

an

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ogra

ptu

s p

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um

us

(Ric

hte

r) N

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iver

sogr

aptu

s sp

.

ex .g

r. N

. nil

son

i (B

arr.

)

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alin

a s

p. a

ff . D

. fi

del

is (

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r.),

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nu

laca

rdiu

m c

f. e

volv

ens

Bar

r.

Car

dio

lita

ex.

gr. C

. boh

emic

a B

arr.

‘Ort

hoc

era

s’ c

f. p

rim

aev

um

(Fo

rbes

)

Par

akio

noc

era

s sp

.

Gei

son

ocer

as

sp.

Mic

hel

inoc

era

s sp

.

Lo

wer

Lu

dlo

w

Bo

reh

ole

s

Tu

zla-

Co

stin

eşti

, M

anga

lia,

Căl

ăraş

i, Ţ

ănd

ărei

and

Bir

uin

ţa

Hol

oret

ioli

tes

(Bal

tico

grap

tus)

bal

ticu

s E

is.

Ple

ctog

rapt

us

ma

cile

ntu

s (T

örn

q.)

Mon

ogra

ptu

s u

nci

nat

us

Tu

llb.

M. i

nci

pien

s (B

arr.

)

Mon

ocli

ma

cis

mic

ropo

ma

(Ja

eck

.)

Pri

stio

grap

tus

du

biu

s (S

ues

s)

Saet

ogra

ptu

s co

llon

us

(Bar

r.)

S. c

him

aer

a (

Tu

llb.

)

Boh

emog

rapt

us

boh

emic

us

(Bar

r.)

Lob

ogra

ptu

s sc

anic

us

Tu

llb.

Neo

div

erso

grap

tus

nil

sson

i (B

arr.

).

Up

per

Wen

lock

Bo

reh

ole

s Ia

nca

-Ber

lesc

u

and

Bo

rdei

Ver

de

Ple

ctog

rapt

us

pra

ema

cile

ntu

s B

ou

č. e

t M

ün

ch

Mon

ogra

ptu

s fl

emin

gii

(Sal

ter)

Mon

ocli

ma

cis

fl u

men

dos

ae

(Go

rt.)

Pri

stio

grap

tus

du

biu

s (S

ues

s)

P. p

seu

dod

ubi

us

(Bo

uč.

)

Cyr

togr

aptu

s tr

ille

ri E

isel

C. l

un

dgr

eni

Tu

llb.

Bu

tovi

cell

a m

igra

ns

(Bar

r.)

Car

dio

la s

p. c

f. C

. in

terr

upt

a S

ow

.

‘Ort

hoc

era

s’ s

p. e

x. g

r.

Pri

ma

evu

m (

Fo

rbes

)

Gei

son

ocer

as

sp.

Mic

hel

inoc

era

s sp

.

Lo

wer

Wen

lock

Bo

reh

ole

s Ţ

ănd

ărei

an

d B

ord

ei

Ver

de

Ret

ioli

tes

gein

itzi

anu

s B

arr.

Bar

ran

deo

grap

tus

pulc

hel

lus

(Tu

llb.

)

Mon

ogra

ptu

s pr

iod

on (

Bro

nn

)

M. fi

rm

us

(Bo

uč.

)

M. p

seu

doc

ult

ellu

s B

ou

č.

M. k

olih

ai B

ou

č.

Mon

ocli

ma

cis

vom

erin

a v

omer

ina

Nic

h.

M. v

omer

ina

ba

sili

ca (

Lap

w.)

M. l

inn

arso

ni

(Tu

llb.

)

Pri

stio

grap

tus

pra

edu

biu

s (B

ou

č.)

P. c

f. d

ubi

us

(Su

ess)

Cyr

togr

aptu

s m

urc

his

oni

Car

r.

C. s

p.1

.

Ord

ovi

cian

Bo

reh

ole

Man

gali

a

Did

ymog

rapt

us

hir

un

do

(Sal

t.)

D. c

f. e

xten

sus

(Hal

l)

Lin

gule

lla

aff

. dav

isii

Mco

y

PALAEOZOIC FROM DOBROGEA AND PRE-DOBROGEA

704

Tab

le 4

. Th

e

mai

n z

on

es a

nd

fau

nas

of

the

Ord

ovi

cian

–E

msi

an d

epo

sits

of

Eas

t M

oes

ia (

aft e

r R

ăile

anu

et

al.

19

66

; P

atru

liu

s &

Io

rdan

19

69

, 1

97

0;

Iord

an e

t a

l. 1

98

5;

Iord

an

19

99

; Vai

da

et a

l. 2

00

4) Z

on

esT

rilo

bit

es /

Ten

tacu

liti

ds

/

Co

no

do

nts

/ C

hit

ino

zoan

sB

rach

iop

od

sP

lan

ts/

Biv

alve

s/F

ora

min

ifer

a

Em

sian

Pse

ud

ocry

pha

eus

pros

tell

ans

Pil

leti

na

asi

atic

a

Dip

leu

ra f

orn

ix

Pse

ud

ocry

pha

eus

pros

tell

ans

pec

tin

ata

(R

oem

)

P. h

amm

ersc

hm

idti

(R

oem

.)

Par

ahom

alon

otu

s ge

rvil

ei r

um

ania

na

Io

rdan

Bu

rmei

ster

ia r

ăile

anu

i Io

rdan

Dip

leu

ra f

orn

ix H

aas

Lep

tost

roph

ia i

nd

ex H

avl.

Sch

izop

hor

ia m

ult

istr

iata

(H

all)

Mu

tati

onel

a p

odol

ica

Ko

zl.

Sch

elw

ien

ella

um

bra

culu

m

(Sch

l.)

Ch

onet

es u

nke

len

sis

Dah

m.

Fim

bris

piri

fer

trig

eri

(Ver

n.)

F. d

alei

den

sis

(Ste

in.)

Naj

ad

ospi

rife

r n

aja

du

m (

Bar

r.)

Cos

tisp

irif

er a

ren

osu

s (C

on

rd.)

Act

inop

teri

a c

osta

ta (

Go

ldf.

)

Lei

opte

ria

sp

.

Par

alle

lod

on m

and

elen

sis

(Dah

m.)

Ort

hon

ota

tri

plic

ata

Fu

chs

Par

acy

cla

s ru

gosa

(G

old

f.)

Car

ydiu

m i

nfl

atu

m D

ien

st.

Cyp

rica

rdin

ia s

p.

Pra

ecte

nod

onta

sp

.

Gra

mm

ysia

man

gali

ca I

ord

an

G. a

ff . a

rmor

ica

Mu

n. C

h.

Gra

mm

ysio

idea

in

aequ

alis

cala

tica

Io

rdan

Nu

culi

tes

elli

ptic

us

(Mau

r)

Pal

eoso

len

cos

tatu

s Sa

nd

b.

Pra

gia

n

Now

akia

acu

aria

Vol

iyn

ites

vel

ain

ii

Ten

tacu

lite

s gy

raca

nth

us

(Eat

on

)

T. s

tra

elen

i M

aill

.

T. o

rnat

us

(So

w.)

Pro

lati

onu

s pr

ael

ongu

s L

jasc

h.

Mu

ltic

onu

s m

aca

rovi

ci I

ord

an C

orn

icu

lin

a s

p.

Lo

chk

ovi

an

Icri

odu

s w

osch

mid

ti

Eis

ena

ckit

ina

boh

emic

a

Uro

chit

ina

sim

plex

Cin

gulo

chit

ina

plu

squ

elle

ci

C. s

erra

ta, C

. erv

ensi

s

Fu

ngo

chit

ina

lata

Mar

gach

itin

a c

aten

aria

Arm

oric

och

itin

a s

p.

An

cyro

chit

ina

sp

.

Pří

do

lí

Up

per

Lu

dlo

w

Lo

wer

Lu

dlo

w

Up

per

Wen

lock

Lo

wer

Wen

lock

ult

imu

s-fo

rmos

us

inex

pec

tatu

s

boh

emic

us

inci

pien

s

nil

sson

i-sc

anic

us

lun

dgr

eni

rad

ian

s

fi rm

us

mu

rch

ison

i

inse

ctu

s-ce

ntr

ifu

gus

Eis

ena

chit

ina

fi l

ifer

a

E. l

agen

omor

pha

Urn

och

itin

a u

rna

(E

isen

ack

)

Urn

och

itin

a k

amel

i B

ou

men

dje

l

Lin

och

itin

a k

lon

ken

sis

Par

is &

Lau

feld

Bu

rsa

chti

na

kri

zi P

aris

an

d L

aufe

ld

Ord

ovi

cian

hir

un

do

Lin

gule

lla

aff

. dav

isii

Mco

y

A. SEGHEDI

705

Obrucev, Drepanaspis sp., ?Kiaeraspis sp.) (Patrulius & Iordan 1969, 1970), along with brachiopods and crinoids are associated with the tentaculite Nowakia maureri Zagora and Homoctenus cf. hanusi (Bouc. et Prantl.) (Iordan 1999) (Table 5).

Th e Givetian is indicated by the tentaculite and brachiopod zones Tentaculites conicus and respectively Fimbrispirifer venustus and Mucrospirifer mucronatus. Th e faunal assemblage (Table 5) includes tentaculites, brachiopods and stromatoporids (Iordan 1967, 1981, 1988; Iordan et al. 1985).

Th e Frasnian is documented by the Homoctenus krestovnikovi tentaculite zone and Mucrospirifer bouchardi brachiopod zone. Th is assemblage (Table 5) includes tentaculites and brachiopods (Iordan 1967, 1981, 1988).

Th e Famennian (Viroaga Limestones), identifi ed only in the Viroaga borehole from South Dobrogea, is indicated by the brachiopod zone Cyrtospirifer archiaci and the conodont zone Palmatolepis crepida (Iordan et al. 1987a); the assemblage consists of brachiopods and conodonts, along with holothurian sclerites, ophiuroids, Moravamminids and calcispheres (Table 5).

Carboniferous – Th e Middle–Upper Visean reaches 268 m thick in the Dobromiru borehole from South Dobrogea. Th e facies succession in this borehole consists of dark limestone intervals up to 11 m thick, interbedded with black and grey mudstones and dark sandstones rich in plant debris (Iordan et al. 1987b; Iordan 1999).

In contrast with the Upper Devonian carbonate succession, the limestones and dolomites of the Lower Carboniferous facies are devoid of evaporites. Th eir maximum thickness is 1580 m, attained in the 2881 Călăraşi borehole. Th e Visean bio- and lithofacies includes endothyracee and algae (Paraschiv & Năstăseanu 1976; Pană 1997); the outer platform facies developed over large areas (Vinogradov & Popescu 1984).

Th e upper part of the Carboniferous consists of a coal-bearing clastic succession about 350 m thick (Vlaşin Formation, Namurian–Westphalian) (Paraschiv & Popescu 1980; Paraschiv et al. 1983), unconformably overlying the Middle Devonian or

Lower Carboniferous carbonate rocks. Th e clastic succession consists of clays, siltstones and sandstones rich in plant debris, with limestone intercalations and coal layers. Th e lithofacies is deltaic in the Bashkirian and sabkha-type in the Moskovian and Bashkirian (Pană 1997). In places, basic and acid volcanic rocks are associated with the clastics (Paraschiv 1986c). Plant remains are abundant (Calamites, Stigmaria, etc.) (Iordan 1999).

Permian – Th e Permian deposits are little represented in East Moesia, despite their large-scale development in West Moesia. Sedimentary breccias from the Dobromiru borehole, 54 m thick, lying unconformably on top of lower Upper Carboniferous clastics and unconformably overlain by Jurassic sandstones, have been ascribed to the Permian (Iordan et al. 1987b). Th e proximal clastic facies of the Permian suggest deposition next to a tectonic uplift , very likely the rift shoulder referred to as the North Bulgarian high. Th e typical Permian facies of continental red-beds is widespread in West Moesia. West of Mangalia, a Permian limestone facies was identifi ed based on foraminifera (Pană 1997). It represents a continuation of the limestone facies of the Upper Carboniferous.

Palaeocontinental Affi nities

During the Neoproterozoic, three accretionary orogens may be defi ned along continental margins (Ziegler 1990). Th e Baikalian-Timanian orogen is confi ned to processes at the edge of the Baltica palaeocontinent (or East European Craton), while the Pan-African and Cadomian are related to Gondwana margin.

Peri-Gondwanan terranes, originating from the northern and western margins of Gondwana, are separated into 4 types, based on their detrital zircon sources: Avalonian, Ganderian, Cadomian and Cratonic (Nance et al. 2008). Avalonian terranes represent oceanic volcanic arcs within the Panthalassa Ocean surrounding the Rodinia supercontinent, accreted to the Gondwana margin by 650 Ma. Th e main source for Avalonian detrital zircon was the Amazonian craton. Th e Ganderian and Cratonic terranes have peri-Amazonian detrital zircon sources,

PALAEOZOIC FROM DOBROGEA AND PRE-DOBROGEA

706

Tab

le 5

. Th

e

mai

n z

on

es a

nd

fau

nas

of

the

Eif

elia

n–

Wes

tph

alia

n d

epo

sits

of

Eas

t M

oes

ia (

aft e

r R

ăile

anu

et

al.

19

66

; Pat

ruli

us

& I

ord

an 1

96

9, 1

97

0; I

ord

an e

t a

l. 1

98

5; I

ord

an

19

99

; Vai

da

et a

l. 2

004;

Kö

nig

sho

f &

Bo

nch

eva

2005

).

Zon

esC

onod

onts

/ Ten

tacu

litid

s/Tr

ilobi

tes/

Chi

tinoz

oans

Bra

chio

pods

/Str

omat

opor

ids

Plan

ts/ B

ival

ves /

For

amin

ifera

/ Pla

code

rm &

O

stra

code

rm F

ishe

s

Nam

uria

n-W

estp

halia

nD

icty

oclo

stus

retif

orm

isPr

oduc

tus

carb

onar

ius

Posi

doni

a co

rrug

ate

Jan

eia

prim

aeva

Cal

amite

sSt

igm

aria

Vis

ean

Gna

thod

us s

ymm

etri

cus

Meg

acho

nete

s zi

mm

erm

ani

Dav

iese

lla ll

ango

lens

is (D

av.)

Dic

tyoc

lost

us s

emire

ticul

atus

D. m

ultis

pini

feru

sG

igan

topr

oduc

tus

giga

nteu

sG

. bis

ati

Avic

ulop

ecte

n in

ters

tria

lis

Endo

thyr

a br

adyi

Tour

nais

ian

Syph

onod

ella

isos

ticha

Qua

sien

doth

yra

cobe

itusa

naSy

phon

odel

la is

ostic

haQ

uasi

endo

thyr

a co

beitu

sana

Fam

enni

anC

yrto

spir

ifer

arch

iaci

Pal

mat

olep

is c

repi

daPo

lygn

athu

s no

doco

stat

us B

rans

. et M

ehl.

Palm

atol

epis

stop

peli

Sand

. Ic

riod

us c

ornu

tus

Sand

.H

ibba

rdel

la c

f. or

tha

Palm

atol

epis

gla

bra

lept

a P.

gra

cilis

gra

cilis

Cyr

tosp

irife

r ar

chia

ci (M

urch

.) C

. tch

erny

chev

i sib

iric

a Iv

ania

C. s

pp.

Prod

ucte

lla sp

. At

hyri

s sp

.M

egac

hone

tes s

p.

Fras

nian

Hom

octe

nus

kres

tovn

ikov

i Muc

rosp

irife

r bo

ucha

rdi

Hom

octe

nus

kres

tovn

ikov

i Lja

sh.

Dic

rico

conu

s sp

. A.

Tent

acul

ites

sp. F

ex.

gr.

T. s

trae

leni

Mai

ll.Po

lygn

athu

s w

ebbi

P. d

ecor

osus

P. d

ubiu

sP.

xyl

us x

ylus

Oza

rkod

ina

cong

esta

Cho

nete

s row

ei C

l. et

Sw

. A

thyr

is v

ittat

a H

avl.

Muc

rosp

irife

r bou

char

di (M

urch

.)M

edio

spir

ifer

auda

cula

(Con

r.)

Spin

ocyr

tia a

ff. m

artia

noffi

(Stu

ck.)

Giv

etia

nTe

ntac

ulite

s co

nicu

s Fi

mbr

ispi

rife

r ve

nust

us

Muc

rosp

irife

r m

ucro

natu

sTe

ntac

ulite

s con

icus

Roe

m.

T. b

ellu

lus

poto

mac

ensi

s (P

ross

.) H

omoc

tenu

s sp

.Ic

riod

us b

revi

sI.

diffi

cilis

Poly

gnat

hus

varc

us,

P. a

nsat

usP.

ling

uifo

rmis

wed

dige

iP.

par

aweb

biP.

xyl

us x

ylus

O

zark

odin

a ba

llai

O. l

ata

Stro

phod

onta

(S.)

dem

issa

(Con

r.)D

evon

ocho

nete

s sc

itulu

s (H

all)

Atry

pa re

ticul

aris

kuz

basi

ca R

zhon

.Fi

mbr

ispi

rife

r ve

nust

us (H

all)

Muc

rosp

irife

r muc

rona

tus

(Con

r.)

Amph

ipor

a ra

mos

a Ph

ill.

Eife

lian

N. m

aure

riN

owak

ia m

aure

ri Z

agor

a H

omoc

tenu

s cf

. han

usi (

Bou

c. e

t Pra

ntl.)

Poly

gnat

hus

lingu

iform

is a

lveo

lus

P. li

ngui

form

is g

amm

a m

orph

.P.

xyl

us x

ylus

Icri

odus

nor

ford

iI.

orri

Pa

nder

odus

exp

ansa

Fim

bris

piri

fer

sp.

Rhip

idom

ella

pen

elop

e (H

all)

Lept

ostro

phia

rotu

nda

Bul

b.

Mar

kito

echi

a m

arki

(Hav

l.)

Pseu

dosp

oroc

honu

s kr

ejci

i Pot

. et B

ern.

Aneu

roph

yton

ger

man

icum

Kr.

et W

eyl.

Cal

amop

hyto

n pr

imae

vum

Kr.

et W

eyl.

Hye

nia

sp.

Luna

spis

bro

ilii G

ross

L. s

p.?

Wijd

easp

is sp

. Obr

ucev

D

repa

nasp

is sp

.?K

iaer

aspi

s sp

.

A. SEGHEDI

707

formed either of recycled crust, or from material not

aff ected by Avalonian-Cadomian continental margin

magmatism. Th e Cadomian terranes were supplied

with detrital zircons derived from the African craton.

Th e confi guration of palaeocontinents at the

end of the Neoproterozoic is shown in Figure 19.

Laurentia, which started to separate from West

Gondwana opening the Iapetus Ocean, includes

North America, Greenland and the British Isles north

of the Iapetus Suture (Cocks & Torsvik 2005). Baltica

includes most of Scandinavia and Russia eastward

to the Urals. Peri-Gondwanan terranes of West and

East Avalonia include British Isles and NW Europe

south of the Iapetus Suture, Eastern Newfoundland,

most of the Maritime Provinces of Canada, parts of

the eastern United States (Cocks & Torsvik 2005).

Terranes forming the basement of Danubian Nappes

in the South Carpathians, as well as the East Moesian

basement and the Boclugea terrane are included here

(Balintoni et al. 2010). Cadomia includes Armorica,

Iberia, Saxothuringia, Tepla-Barrandia, proto-Alps,

Sakarya, Eastern Pontides, Menderes and Kırşehir

(Linnemann et al. 2008) to which the Carpathian and

Orliga terranes were added (Balintoni et al. 2010).

Th e Scythian Platform – Considering the two main

models proposed for its evolution, the Scythian

Platform may show either Baltican or Avalonian/

Cadomian affi nities. Discrimination between the two

models is not possible yet based on detrital zircon

data, which are still lacking.

If the Scythian Platform represents the southern

margin of the East European Craton, involved in

Neoproterozoic deformation, it was part of the

Baltica palaeocontinent. Evidence in favor of this

affi nity comes from stratigraphical and geophysical

data. As with the Neoproterozoic–Lower Palaeozoic

cover of the East European Platform, the undeformed

platform cover of the Scythian Platform starts with

the Neoproterozoic clastics and contains Silurian

carbonates. Geophysical evidence also indicates

that the lithosphere and crust of the southern East

European Platform and Scythian Platform are

thicker than those of the Phanerozoic Western

Europe terranes (Yegorova et al. 2004; Yegorova &

Starostenko 2002).

According to the ‘Scythian orogen’ model, the southernmost part of the East European Platform formed a passive margin of Baltica from the Cambrian until the collision and docking of the Scythian Platform crust in the Late Palaeozoic, which in this case would include terranes with both Baltican and peri-Gondwanan affi nity. Th is model is based on the assumption that the Scythian Platform represents an eastward prolongation of the Variscan orogen (or Palaeozoic platform) of central and western Europe (Mouratov & Tseisler 1982; Milanovsky 1987; Zonenshain et al. 1990; Visarion et al. 1993; Nikishin et al. 1996, 2001). However, there is no irrevocable evidence that the basement crust of the Scythian Platform includes a Late Palaeozoic (Variscan) accretionary orogenic belt (Seghedi et al. 2003; Stephenson et al. 2004). As observed in borehole cores, the Palaeozoic successions of the Scythian Platform do not show a penetrative deformation related to folding (Seghedi et al. 2003b), as seen in North Dobrogea. More steeply inclined dips of bedding noticed in boreholes can be explained as result of block tilting and rotation connected to Permian rift ing (Seghedi et al. 2003b; Saintot et al. 2006).

North Dobrogea – Results from detrital zircons from 5 samples of Boclugea lithologies indicate several age groups, corresponding to the late Neoproterozoic–early Cambrian, the Mesoproterozoic, the 1.7 and 2.2 Ga interval of the Palaeoproterozoic and the 2.55–2.85 Ga interval of the Neoarchaean, with age gaps or minima in the intervals 0.75–0.85 Ga, 1.65–1.7 Ga and 2.2–2.5 Ga (Balintoni et al. 2010). Considering the signifi cant input from Mesoproterozoic and Palaeoproterozoic sources and only minor zircon contributions from Grenvillian sources, the authors conclude that an East Avalonian palaeogeographic affi nity seems likely for the Boclugea terrane.

Detrital zircon ages for the Orliga terrane show clusters at 0.5–0.75 Ga, 1.75–2.1 Ga, 2.3–2.5 Ga and 2.55–2.85 Ga, with a gap covering the entire Mesoproterozoic and the Grenvillian (Balintoni et al. 2010). Because such patterns characterize the Cadomian terranes that originated close to the West African craton (Samson et al. 2005; Nance & Murphy 1994), the Orliga terrane is considered a true Cadomian terrane (Balintoni et al. 2010).

PALAEOZOIC FROM DOBROGEA AND PRE-DOBROGEA

708

Based on U-Pb detrital zircon ages, a correlation of Orliga with the Cadomian Sebeş-Lotru terrane and of Boclugea with the Avalonian Danubian terranes from the South Carpathians is possible (Balintoni et al. 2010). Th is correlation suggests that the South Carpathians and North Dobrogea were connected during the Variscan orogeny. Th e continuity was disrupted in the Early Cretaceous, by westward displacement of Moesia, related to the opening of the western Black Sea basin (Balintoni & Baier 1997).

East Moesia

Evidence for the palaeogeographic affi nity of East Moesia is still controversial. Based on a detailed analysis of tectonostratigraphic, palaeontological and geochronological data, four distinct terranes,

two with Baltican and two with Avalonian affi nities have been separated in the entire Moesian Platform (Oczlon et al. 2007). According to these authors, East Moesia consists of an Avalonian terrane comprising Central and a large part of South Dobrogea, separated by the narrow Palazu terrane with a basement of Baltican affi nity. Th e review of terranes along the southwestern margin of the East European Craton, between the North Sea and the Black Sea, suggests that dextral strike-slip dominated the southwestern Baltican margin during Late Ordovician–Early Devonian accretion of Far Eastern Avalonia (Oczlon et al. 2007). Southward displacement of some Avalonian terranes, including East Moesia, occured consequently to the Variscan indentation of the Bohemian Massif, while current juxtaposition of the Moesian terranes is the result of sinistral strike-slip

Figure 19. Reconstruction of the Late Vendian (ca. 550 Ma) (modifi ed from Cocks & Torsvik 2005). Spreading centres are

shown as black lines, subduction zones are shown as lines with ticks. Transform faults are accompanied by arrows.

Black areas represent Grenville-Sveconorwegian mobile belts (about 1 Ga). Dark grey shaded areas are regions

that show evidence of Avalonian-Cadomian-Pan-African-Baikalian-Timanian deformation, terrane amalgamation

and magmatism. East Moesia (EM), Boclugea, Danubian, and other correlatives in the Balkans (DB) represent

East Avalonian Peri-Gondwanan terranes (Balintoni et al. 2010). Cadomia (fi gured aft er Linnemann et al. 2008),

includes Armorica (A), Iberia (Ib), Saxothuringia (ST), Tepla-Barrandia (TB), proto-Alps (PA), Sakarya, Eastern

Pontides, Menderes and Kirshehir, with the Carpathian terranes (C) and Orliga (O) added by Balintoni et al. (2010).

P represents Palazu terrane.

A. SEGHEDI

709

displacement during the opening of Mesozoic back-arc basins.

Palynological evidence based on Chitinozoans show a North Gondwanan affi nity of the East Moesian Lower Palaeozoic deposits (Vaida et al. 2004; Vaida & Verniers 2005), in good agreement with the affi nity of the Lochkovian chitinozoans from West Moesia (Northwest Bulgaria) (Lakova 1995; Yanev et al. 2005). Th is however contrasts with the results from miospores (Steemans & Lakova 2004), which support a Baltican affi nity for Moesia at that time.

Lithology and macrofauna of the Lower Palaeozoic in the 5082 Mangalia borehole were previously used to suggest an affi nity with the macrofauna identifi ed in the Bosfor area (Bithynia), in basins with continuous marine sedimentation across the Silurian/Devonian boundary (Iordan 1967). Th e macrofauna of trilobites and tentaculites is also comparable to that from the Armorican and Bohemian Massifs that are considered part of Cadomia (Linnemann et al. 2008).

Detrital zircon data exist so far for the Ediacaran basement of East Moesia. Detrital zircons yielded U/Pb SHRIMP ages of 1497±8 Ma, 1050±1 Ma, 603±5 Ma and 579±7 Ma (Żelaźniewicz et al. 2001), interpreted as Avalonian (Seghedi et al. 2005a, b) or Far East Avalonian sources (Oczlon et al. 2007).

Detrital zircons published by Balintoni et al. (2011) show the following age concentrations: the greatest age grouping of 2.65–3.0 Ga is Archean; Palaeoproterozoic ages show a very low frequency around 1.8 Ga and between 2.2–2.6 Ga; the Mesoproterozoic ages peak around 1.5 Ga; Palaeoproterozoic ages cluster around 1.7 Ga and between 1.95–2.2 Ga; late Neoproterozoic ages between 633–583 Ma., suggesting that the Brasiliano orgen (Nance et al. 2009) was the most important Neoproterozoic detrital zircon source for the Histria Formation. A detailed discussion of the zircon sources and their discrimination is found in Balintoni et al. (2011).

Discussion

Th e geological evolution of the Palaeozoic terranes from Dobrogea and Pre-Dobrogea took place during several events occurring at the south-western margin of the East European Craton: Early Palaeozoic

accretion of peri-Gondwanan terranes of Far East Avalonia (Pharaoh 1999; Winchester et al. 2002, 2006); accretion to Laurussia during the Hercynian orogeny; rift ing and displacement along the TESZ, together with slivers of proximal Baltican terranes accommodated along strike-slip faults (Winchester et al. 2006; Oczlon et al. 2007).

Palaeogeographic affi nities proposed for the terranes presented here are based mainly on lithology and palaeontological evidence derived from the Palaeozoic record. Detrital zircon ages are published only for the Central Dobrogea basement (Żelaźniewicz et al. 2001, 2009; Balintoni et al. 2011) and for the early Early Palaeozoic Boclugea and Orliga terranes later aff ected by Hercynian regional metamorphism (Balintoni et al. 2010) (Figure 8). Th e existing evidence indicates that terranes with Baltican, Avalonian and Cadomian affi nities occur in the Dobrogea and Pre-Dobrogea area.

Baltican affi nities are shown by the Palazu terrane and the Pre-Dobrogea basement. Th e Palazu terrane, isolated in the basement of South Dobrogea between the Capidava-Ovidiu and Palazu faults, is a proximal Baltican sliver (Oczlon et al. 2007). It is made of Archaean and Palaeoproterozoic crust, similar in composition to the Ukrainian shield (Giuşcă et al. 1967, 1976), or the Sarmatia segment of Baltica (as defi ned by Bogdanova et al. 1997). Ediacaran rift -related volcano-sedimentary successions (Cocoşu Group) are preserved on top of the metamorphic crust. Th is geological association suggests an initial location along the TESZ margin near the Volhyn-Orsha aulacogen, about 500 km north of the present location of the Palazu terrane. Th e Volhynian series from the Volhyn depression developed within the aulacogen includes lower Vendian fl ood basalts formed during continental rift ing that accompanied Rodinia breakup at about 750 Ma (Kumpulainen & Nystuen 1985; Torsvik et al. 1996; Nikishin et al. 1996). Lithological, petrographic and geochemical features of the Volhynian basalts (Bakun-Czubarow et al. 2002; Białowolska et al. 2002) correlate well with those of the Cocoşu Group (Seghedi et al. 2000).

In a series of palaeogeographic maps for the East European Craton (Nikishin et al. 1996), the Pre-Dobrogea area is considered part of the Baltica palaeocontinent. Fine-grained Vendian lithologies

PALAEOZOIC FROM DOBROGEA AND PRE-DOBROGEA

710

from Pre-Dobrogea, with phosphate nodules and

Vendotaenia antiqua, accumulated in the Dnestr

basin of Baltica. Th is is one of the very large shelf

basins, established along the western slope of the East

European Craton during the Ediacaran and referred

to as the Peri-Tornquist basin (Nikishin et al. 1996).

During the Cambrian, the Dnestr basin developed

into a passive continental margin, with shallow-

marine clastics overstepping the Vendian deposits.

Th is evolution is confi rmed in the Aluat and Bârlad

basins, where Cambrian clastics have been identifi ed.

Th e typical facies of Silurian carbonates from the

Moldavian and Baltic basins (Sliaupa et al. 2006) was

not found in Pre-Dobrogea. However, the Ordovician

and Silurian deposits, identifi ed by chitinozoans

in the upper part of what had been considered a

Vendian succession, show a siltstone-mudstone

facies with sandstone and limestone interbeds (Vaida

& Seghedi 1997). Th is suggests that the transition

from the carbonate facies of the Peri-Tornquist Basin

to the fi ne-grained clastic facies of the Peri-Caspian

basin (Nikishin et al. 1996) took place by lateral facies

changes in the Pre-Dobrogea area.

Th e Avalonian affi nity of East Moesian

Lower Palaeozoic successions is suggested by the

unconformities within the Silurian and at the

base of the Silurian and Devonian, characteristic

of Caledonian displaced terranes along the TESZ

(Oczlon et al. 2007). Th e peri-Gondwanan, Avalonian

provenance of Central Dobrogea is based on the

detrital zircon spectrum in the Ediacaran turbidites

presented so far (Żelaźniewicz et al. 2001, 2009;

Balintoni et al. 2011) and on cold water acritarchs

from the East Moesian Palaeozoic cover (Oczlon et

al. 2007).

In North Dobrogea, terranes with both Avalonian

and Cadomian affi nities are indicated by detrital

zircon ages (Balintoni et al. 2010). Th e location

of the Palaeozoic shelf facies south of deep marine

Palaeozoic sediments indicates a north-northeast

facing continental margin of North Dobrogea,

consistent with a North Gondwanan provenance

(Oczlon et al. 2007). Th e Măcin-type Palaeozoic

successions overlie the Cambrian clastics of the

Avalonian Boclugea terrane, so they might have

travelled together across the Tornquist sea. Initial

relations, as well as the progressive metamorphism

from the low-grade facies Boclugea rocks to the very low-grade Silurian –lower Devonian deposits is obscured by subsequent thrusting, due to involvement in the Hercynian orogeny, following the accretion of the Cadomian Orliga terrane.

Indirect evidence for an Avalonian provenance of the Lower Palaeozoic terranes in North Dobrogea can be derived from a critical analysis of their lithostratigraphy and from comparison with the Palaeozoic successions of East Moesia.

In the Măcin zone of North Dobrogea, the basinal, anoxic Silurian lithofacies is shallowing upward to Lower Devonian shelf sandstones and carbonates (Figure 11). Th e biostratigraphy of these successions is ill-defi ned due to scarcity of fossil remains, the only exception being the Bujoare fossil site rich in Lower Devonian macrofauna. Th e entire Middle–Upper Devonian section is missing here due to erosion. Th is is indicated by the abundant limestone clasts reworked in the lowest fanglomerates of the Carapelit Formation exposed near the Lower Devonian outcrops. As the Carapelit conglomerates are proximal deposits, provenance from more distal sources of conodont-bearing limestone clasts is precluded. Clast petrography and their fossil content in the lower members of the Carapelit Formation represent evidence that the middle and late Devonian in the Măcin zone developed in limestone facies, similar to East Moesia and the Scythian Platform.

Considering all lithological and structural information about the Silurian–Devonian successions in the Măcin zone of North Dobrogea, lithological correlation with East Moesian coeval deposits leads to a more reasonable lithostratigraphy, even if palaeontological evidence is scarce. Th e dark shales of the Cerna Formation might be ascribed to the ?Wenlock–Lower Ludlow, based on their black shale facies. Th e overlying black argillites might represent the Upper Ludlow, as recorded in lithological logs of several boreholes from East Moesia, where the transition to upper Ludlow is marked by a change in lithology from dark shales to argillites (Iordan 1999; Seghedi et al. 2005a). Th e overlying argillites with thin interbeds of sandstones and calcarenites could represent the Pridoli. Discontinuous, thin layers of volcanic tuff s are interbedded in the Upper Ludlow–Pridoli succession in the Măcin zone, as with the Pridoli

A. SEGHEDI

711

lithofacies from the Zăvoaia borehole in East Moesia. Identifi cation of the conodont Icriodus woschmidti in the same lithofacies marks the presence of the Lower Devonian, indicating the continuity of the argillite facies into the Lochkovian. Th e Lochkovian, well-dated on the brachiopod-dominated macrofauna from the Bujoare Hills, contains several beds of mica-rich quartzitic sandstones, capped by a thin orthoquartzite member, which might be lithologically correlated with the Eifelian lithofacies of East Moesia. Th e shallow-marine limestones overlying the Eifelian quartzites, rich in crinoidal ossicles and silicifi ed solitary corals, could be ascribed to the Givetian. Th e rest of the Devonian shallow marine carbonate facies succession was probably removed by erosion during deposition of the Carapelit Formation.

Following this reasoning, it is possible to conclude that the Măcin-type Silurian–Devonian deposits of North Dobrogea represent East Moesian successions. Based on a previous palaeotectonic model (Mosar & Seghedi 1999) (Figure 20) it is proposed that during the Cambrian–Silurian, the Avalonian part of North Dobrogea was part of the East Moesian terrane, docked to Baltica during the Silurian, and experiencing a common history with the Scythian margin of Baltica during the Devonian. Th e terranes were further involved in the complex evolution of the Hercynian orogen, when the Avalonian North Dobrogea represented the upper plate where regional metamorphism, crustal melting and granitic plutonism and volcanism were taking place. Th e Cadomian Orliga terrane became part of North Dobrogea in the Late Carboniferous, aft er the closure of the Rheic Ocean.

Synthesis and Conclusions

As part of the Baltica palaeocontinent, the area of the Scythian Platform was a shelf basin from the Vendian until the Silurian. Early Devonian rift ing disrupted the shelf basins that evolved further as a carbonate ramp during the Middle Late Devonian–Lower Carboniferous. Following the Middle–Late Carboniferous subduction in the areas of the West European Variscan and South European Scythian orogens (as defi ned by Ziegler 1989), siliciclastic deposits with coal measures prograded on the passive margin carbonates. Th e sedimentary record suggests

that during the Palaeozoic, the basement of the Pre-Dobrogea Depression was part of the Baltican foreland involved in the amalgamation of Avalonia/Baltica with Laurentia, and further evolved as part of the Variscan foreland.

Palaeozoic deposits of East Moesia show a well-preserved palaeontological record and do not show any structural elements indicative of metamorphism. Th e succession of Palaeozoic facies is well documented based on biostratigraphic zones: graptolites for the Ordovician and Silurian, conodonts and chitinozoans for the Lochkovian, tentaculites and trilobites for the Pragian–early Givetian and spiriferids, conodonts and foraminifera for the Givetian–Visean. Th e Namurian–Westphalian is dated on placoderm and ostracoderm fi shes, terrestrial plants, algae and foraminifera. Gaps in the succession correspond to the Llandovery, Upper Wenlock and lower Upper Ludlow. Based on these unconformities, a peri-Gondwanan Avalonian affi nity is inferred for East Moesia, considering also the Avalonian affi nity of the East Moesian Ediacaran basement.

Th e Precambrian basement of the Moesian Platform between the Capidava-Ovidiu and Palazu faults is a terrane with proximal Baltican derivation. Th is terrane, devoid of Palaeozoic deposits, was probably detached from the northwestern slope of the Ukrainian shield (where Volhyn fl ood basalts overlie the Palaeoproterozoic and Archaean Sarmatia margin), and was displaced along the TESZ.

Ordovician to Devonian deep marine sediments developed south of the Sfantu Gheorghe Fault (Figure 10) show a strong contrast in facies and structural style with the Lower Palaeozoic passive margin successions of the Scythian Platform. Th e Tulcea-type Palaeozoic sediments, dated on conodonts and palynological assemblages, form several northward-younging, upward-coarsening terranes, consisting of radiolarian cherts, siliceous shales and distal turbidites. Rocks show a southward increase in metamorphic grade, from subgreenschist to lower greenschist facies. Trace fossil assemblages and facies association indicate that turbidites accumulated in deep basins below the CCD. Lithological association and upward-coarsening facies architecture from siliceous pelagic rocks to turbidites suggests that these terranes represent oceanic and trench sediments,

PALAEOZOIC FROM DOBROGEA AND PRE-DOBROGEA

712

Figure 20. Plate tectonic reconstruction showing a possible scenario for the evolution of North Dobrogea, East Moesia and Pre-

Dobrogea and surrounding oceanic realms in the Palaeozoic (modifi ed from Mosar & Seghedi 1999). Th e model is

based on the constitution of North Dobrogea Palaeozoic basement, that includes an Ordovician–Devonian accretionary

wedge separated from a Late Palaeozoic magmatic arc and retro-arc foreland basin by a strike-slip fault. Th e model shows

accretion to Baltica of East Moesia (including the Boclugea terrane), following the closure of the Tornquist Sea and Variscan

accretion of Orliga terrane upon closure of the Rheic Ocean. North Dobrogea shows south-dipping subduction in the Lower

Palaeozoic and fi nally rather gently docks with the SE edge of Baltica and develops a strong syn- to post-collisional strike-

slip deformation along the Teyssiere zone. Th is late Hercynian wrench tectonics is also related to the destruction of the

Variscan Orogen and strongly aff ects the Dobrogea accretionary prism remnants, as well as the continental margin arc.

A. SEGHEDI

713

which record the trenchward movement of an oceanic plate. Th ey are interpreted as underthrust units of an accretionary wedge, accreted to the upper plate during southward subduction of the Scythian margin of the East European Craton (Mosar & Seghedi 1998; Seghedi 1999).

Th e Upper Palaeozoic Carapelit Formation, ascribed to the Upper Carboniferous–Early Permian, obviously represents a continental, volcanic molasse, intruded by calc-alkaline granites. It contrasts in composition and style with the Carboniferous coal measures from the Scythian Platform of Pre-Dobrogea. Th e Carapelit Formation strongly contrasts in facies and explosive calc-alkaline volcanism with the Lower Permian continental red bed facies from the Pre-Dobrogea Depression, accompanied by eff usive-explosive alkaline bimodal volcanism. A model explaining the Upper Palaeozoic continental sedimentation and magmatism in North Dobrogea as a result of thrust propagation in a retro-arc foreland basin (Seghedi et al. 1987; Oaie & Seghedi 1994; Seghedi & Oaie 1995) (Figure 20c) accounts for: the great thickness of fl uvial and alluvial continental sediments of the Carapelit Formation, accumulated in active tectonic condition; mixed clast composition, suggesting sediment recycling; very-low-grade metamorphism (anchizone metamorphism) of the Palaeozoic sediments of the upper plate, connected to retro-arc thrusting and progressive increase in metamorphic grade of Palaeozoic sediments toward the thrusts footwalls.

Compared to East Moesia, the Palaeozoic of North Dobrogea is loosely dated, due to the scarcity of its fossil record. Th is can be explained by the penetrative Hercynian deformation in amphibolite to greenschist and subgreenschist facies conditions and thermal metamorphism connected to granite intrusion. Nevertheless, based on a vertical lithofacies succession, the shallow-marine Lower Palaeozoic of North Dobrogea correlates well with coeval East Moesian deposits. On this basis, an Avalonian affi nity can be inferred for the Măcin-type Lower Palaeozoic, consistent with the Avalonian affi nity of its basement, the Cambrian Boclugea terrane, as suggested by detrital zircons. Unlike East Moesia and Pre-Dobrogea, the continental Late Palaeozoic of North Dobrogea was involved in Hercynian collision, with metamorphism, volcano-plutonic development

and thrusting. Th is suggests an active margin setting and upper plate position. Th e deep marine Palaeozoic pelagic cherts and distal turbidites, were part of the forearc wedge, tectonically accreted to the upper plate along a south-dipping subduction plane. Th e south polarity of the subduction zone is indicated by the deep marine Palaeozoic of North Dobrogea facing the shallow-marine Devonian carbonate platform facies of the Scythian Platform. A north-dipping subduction zone of Palaeotethys developed in the Late Palaeozoic as shown in Figure 20a, b. Th e trace of the Rheic suture runs through North Dobrogea, where both Avalonian and Cadomian terranes are found, as in the Hercynian basement reworked in the Alpine nappes of the South Carpathians (Balintoni et al. 2010).

Permian rift ing accompanied by bimodal alkaline volcanism occurred both in North Dobrogea and the Pre-Dobrogea. In North Dobrogea, Late Permian rift ing is explained as a consequence of extensional collapse of the Variscan orogen, overthickenned by thrusting and granite intrusion (Figure 20d). Th e transition from early Permian syn-collisional calc-alkaline magmatism to Late Permian alkaline magmatism refl ects the changing stress regime from compression to transtension and was explained as result of slab roll-back (Figure 20d). Th e extensional (transtensional) regime propagated from the Hercynian belt into its foreland, where rift ing was accompanied by accumulation of thick volcano-sedimentary successions in the Sărata-Tuzla basin from the eastern part of the rift . Active rift ing of the basement of the Pre-Dobrogea is suggested by emplacement of the bimodal alkaline volcanism of the basalt-trachyte association. Th is volcanism was fed by the suite of the bimodal basalt-trachyte dykes emplaced during the Permian along the northern margin of North Dobrogea, as well as by small bodies of syenites and ultrapotassic alkaline rocks emplaced in the basement highs of the rift basin. In the meantime, alkaline volcanism developed from Late Permian to early Triassic times along the southern border of North Dobrogea, parallel to the Peceneaga-Camena Fault.

Acknowledgements

Th is paper was written as result of the participation in the second ISGB conference in held in Ankara in

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2009 and supported from project PROMED, contract 31-030/2007, fi nanced by the Ministry of Education and Research of Romania. Th e author is extremely grateful to Aral Okay for the strong support and encouragement for writing this paper, as well as for his infi nite patience. Th e comments and suggestions

of Ayda Petek Üstaömer and an unknown reviewer are highly appreciated, considerably increasing the quality of this paper. Th e English language of the paper was greatly improved, thanks to Randell Stephenson.

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