Palaeozoic Formations from Dobrogea and Pre-Dobrogea – An...
Transcript of Palaeozoic Formations from Dobrogea and Pre-Dobrogea – An...
A. SEGHEDI
669
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: [email protected])
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
PALAEOZOIC FROM DOBROGEA AND PRE-DOBROGEA
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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
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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
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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
ure
7.
In
terp
reta
tive
geo
logi
cal
sect
ion
s in
No
rth
Do
bro
gea
sho
win
g th
e H
ercy
nia
n b
asem
ent
and
th
e C
imm
eria
n s
tru
ctu
res
(mo
difi
ed
fro
m S
egh
edi
20
01
). L
oca
tio
n o
f
sect
ion
s is
sh
ow
n o
n in
set
map
. Th
e se
ctio
ns
are
dra
wn
bas
ed o
n o
utc
rop
s an
d b
ore
ho
le d
ata.
Mes
ozo
ic s
tru
ctu
res
are
fi gu
red
aft
er B
altr
es 1
99
3. M
IF–
Mei
dan
chio
i-
Iuli
a F
ault
; T
F–
Tel
iţa
Fau
lt;
K-T
– C
reta
ceo
us–
Ter
tiar
y;
K2–
Lat
e C
reta
ceo
us
‘Bab
adag
bas
in’;
J– J
ura
ssic
; T
– T
rias
sic;
Pz–
un
diff
ere
nti
ated
Her
cyn
ian
bas
emen
t
incl
ud
ing
the
Bo
clu
gea,
Meg
ina
and
Orl
iga
met
amo
rph
ic t
erra
nes
an
d S
ilu
rian
–L
ow
er D
evo
nia
n s
edim
ents
; P–
Per
mia
n a
lkal
ine
com
ple
xes;
C-P
– C
arb
on
ifer
ou
s–
Per
mia
n g
ran
ito
ids;
C2-P
1–
Car
bo
nif
ero
us–
Ear
ly P
erm
ian
Car
apel
it F
orm
atio
n; C
– C
arb
on
ifer
ou
s in
tru
sive
s, C
?– s
up
po
sed
Car
bo
nif
ero
us
cher
ts a
nd
tu
rbid
ites
of
the
Tu
lcea
typ
e P
alae
ozo
ic; E
– E
dia
cara
n t
urb
idit
es o
f C
entr
al D
ob
roge
a; N
P–
Neo
pro
tero
zoic
bas
emen
t o
f C
entr
al D
ob
roge
a; P
DD
– P
red
ob
roge
a D
epre
ssio
n.
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
.
Cri
no
ids
Cte
noc
rin
us
typ
ous
bro
nn
Tec
hn
ocri
nu
s st
riat
us
(Hal
l)
Lop
hoc
rin
us
sp.
En
troc
hu
s sp
.
Co
no
do
nts
Icri
odu
s w
osch
mid
ti
Bry
ozo
ans
Eri
dot
ryp
a s
p.
Fen
este
lla
sp
.
Hem
itry
pa
sp
.
Tri
lob
ites
Pil
leti
na
cf.
ham
mer
sch
mid
ti (
Ric
ht)
Hom
alo
not
us
sp.
Ph
aco
ps s
p.
Sil
uri
anC
ern
a
Fo
rmat
ion
Cya
thop
hyl
lum
Co
no
do
nts
Pan
der
odu
s sp
.
Gra
pto
lite
s R
ast
rite
su
nid
enti
fi ed
gas
tro
po
d
Ord
ovi
cian
Cam
bri
an
Bo
clu
gea
Qu
artz
ites
(Iu
lia
bo
reh
ole
)
Pal
yno
mo
rph
s
Mic
rhys
trid
ium
sp
.
M. l
anat
um
Lei
omar
gin
ata
sim
plex
Gra
nom
argi
nat
a p
rim
a
Un
ipor
ata
nid
ius
Ta
sman
ites
sp
.
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.
‘Ort
hoc
era
s’ v
erte
brat
um
(B
arr.
)
Gei
son
ocer
as
riva
le (
Bar
r.)
Orb
icu
loid
aea
sp
. Lep
tost
roph
ia s
p.
Ple
ctod
onta
sp
.
How
elle
lla
sp
.
Cro
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
Lin
ogra
ptu
s p
osth
um
us
(Ric
hte
r) N
eod
iver
sogr
aptu
s sp
.
ex .g
r. N
. nil
son
i (B
arr.
)
Du
alin
a s
p. a
ff . D
. fi
del
is (
Bar
r.),
Lu
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
PALAEOZOIC FROM DOBROGEA AND PRE-DOBROGEA
714
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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