MEMOIRES DU MUSEUM NATIONAL D'HISTOIRE NATURELLE

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Cover photograph / Photographie de couverture:

The Camian Sidi Stout unconformity in the Dahar cliff (Saharan platform, southern Tunisia). The late Camian Rehach dolomites

rest unconformably over the whole Late Permian to middle Camian sequence. This unconformity is related to middle-late Camian dextral

transcurrent movements along E-W trending faults (see Bouaziz et al. , this volume).

Discordance camienne de Sidi Stour dans la falaise du Dahar (plate-forme saharienne, Tunisie meridionale). Les dolomies de

Rehach (Camien superieur) reposent en discordance sur la serie d'age permien superieur a carnien moyen. Cette discordance resulte du

jeu de grands decrochements dextres orientes est-ouest au Camien moyen-superieur ( Bouaziz et al.. ce volume).

The Peri-Tethys Program, sponsored by international industries (AGIP, ARCO, BRGM, CHEVRON,

CONOCO, EAP(ELF), EXXON, SHELL. SONATRACH, TOTAL), research centres (CNRS, IFP), and a

university (UPMC), started in 1993. It examines the influence of Tethyan evolution on the bordering cratons

since its birth (through the break-up of Pangea), its life (by the extension and formation of oceanic seaways) and

finally its death (by collision between the main bordering plates which led to inversion within the epicratonic

basins).

Volumes deja parus/ Previously published volumes :

Peri-Tethys Memoir 1 (1994): Peri-Tethyan platforms. Proceedings of the IFP/Peri-Tethys Research Conference.

Technip, Paris: 1-275. ISBN: 2-7108-0679-7.

Peri-Tethys Memoir 2 (1996): Structure and Prospects of Alpine Basins and Forelands. Mem. Mus. natn. Hist,

nat ., 170: 1-550 (+ Atlas). ISBN: 2-85653-507-0.

Peri-Tethys Memoir 3 (1998): Stratigraphy and Evolution of Peri-Tethyan Platforms. Mem. Mus. natn. Hist,

nat. , 177: 1-262. ISBN: 2-85653-512-7.

Aussi /Also

Peri-Tethys: stratigraphic correlations 1 (1997). 330 pp. Geodiversitas, 19 (2): 169-499. ISSN: 1280-9659.

Source: MNHN. Paris

Peri-Tethys Memoir 3

Stratigraphy and Evolution

of Peri-Tethyan Platforms

Source: MNHN. Paris

ISBN: 2-85653-512-7

ISSN : 1243-4442

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MEMOIRES DU MUSEUM NATIONAL D'HISTOIRE NATURELLE

TOME 177

Peri-Tethys Memoir 3

Stratigraphy and Evolution

of Peri-Tethyan Platforms

edited by

Sylvie Crasquin-Soleau , n & Eric Barrier' 21

"’Universite Pierre et Marie Curie / CNRS

Departement de Geologie Sedimentaire

T15-25, E.4, case 104

4, Place Jussieu

F-75252 Paris Cedex 05

,2) Universite Pierre et Marie Curie / CNRS

Departement de G6otectonique

T25-26, E.l, case 129

4, Place Jussieu

F-75252 Paris Cedex 05

EDITIONS

DU MUSEUM

PARIS

1998

Source: MNHN. Paris

CONTENTS

SOMMAIRE

Pages

Northern Platform

1. Stratigraphy and sequence stratigraphy of the Upper Carboniferous and Lower

Permian in the Donets Basin. 9

Alain IZART, Celine BRIAND, Denis VASLET, Daniel VaCHARD, Jean BROUTIN, Robert COQUEL,

Alexander MASLO, Natalia MASLO & Raissa KOZITSKAYA

2. Stratigraphic correlations of the Upper Permian and Triassic beds from the

Volga-Ural and Cis-Caspian regions. 35

Edwar A. MOLOSTOVSKY, Ija I. MOLOSTOVSKAYA & Maxim G. MlNIKH

3. The Domanikoid facies of the Russian Platform and basin paleogeography.45

Valentina S. VISHNEVSKAYA

4. Paleomagnetism of Permian to Jurassic formations from the Turan Plate.71

Marie M. LEMAIRE, Evgueni L. GUREVICH, Khodjamourad Nazarov,

Michel WESTPHAL, Hugues FEINBERG & Jean-Pierre POZZI

5. Reconstruction of paleostress fields in Crimea and the North West Caucasus,

relationship with major structures.89

Aline SAINTOT, Jacques ANGELIER, Alexander ILYIN & Oleg GOUSHTCHENKO

6. Neogene evolution of he Carpathian foothills: insights from the

romanian diapir fold area...113

Jean-Claude HlPPOLYTE & Mircea SANDULESCU

7. The Moesian Platform as a key for understanding the geodynamical evolution

of the Carpatho-Balkan alpine system. 127

Fran?oise BERGERAT, Pierre MARTIN & Dimo DlMOV

8 . Scythian Platform: chronostratigraphy and polyphase stages of tectonic history.151

Anatoly M. NlKISHJN, Sierd CLOETINGH, Serguei N. BOLOTOV, Evgenij Yu.

BARABOSCHKIN, Ludmila F. KOPEAVICH, Bronislav P. NAZAREVICH,

Dmitri I. PANOV. Marie-Fran^oise BRUNET, Andrei V. ERSHOV,

Vera V. Ilina, Svetlana S. Kosova & Randell A. STEPHENSON

9. Scythian Platform, Caucasus and Black Sea region: Mesozoic-Cenozoic

tectonic history and dynamics.163

Anatoly M. NlKISHIN, Sierd CLOETINGH, Marie-Fran^oise BRUNET,

Randell A. STEPHENSON, Serguei N. BOLOTOV & Andrei V. ERSHOV

Source: MNHN , Pahs

PERI-TETHYS MEMOIR 3: STRATIGRAPHY AND EVOLUTION OF PERI-TETHYAN PLATFORMS

Southern Platform

10. Triassic series on the Saharan Platform in Algeria; Peri-Tethyan onlaps

and related structuration...177

Hamid Ait SALEM, Sylvie BOURQUIN, Louis COUREL, Berrached FEKIRINE,

Cherfi HELLAL, Leila Mami & Mohamed TEFIANI

11. New data on the Jurassic and Neogene to Quaternary sedimentation

in the Danakil Horst and Northern Afar Depression, Eritrea.193

Mario SAGRI, Ernesto ABBATE, Augusto AZZAROLI, Maria Laura BALESTRIERI,

Marco BENVENUTI, Piero BRUN1, Milvio FAZZUOLI, Giovanni FlCCARELLI,

Marta Marcucci, Mauro Papini, Guilio Pavia, Viviano Reale, Lorenzo Rook

& Tewelde Medhin TECLE

12. Tectonic evolution of the Southern Tethyan margin in Southern Tunisia.215

Samir BOUASSIZ, Eric BARRIER, Jacques ANGELIER, Pierre TRICART

& Mohammed M. TURKI

13. Structural inheritance and kinematics of folding and thrusting along

the front of the Eastern Atlas Mountains (Algeria and Tunisia).237

Dominique FRIZON DE LAMOTTE, Eric MERCIER, Fatima OUTTANI,

Belkacem ADDOUM, Hacene. GHANDRICHE, Jamel OUALI, Samir BOUAZIZ

& Jean ANDRIEUX

Index

253

Stratigraphy and sequence stratigraphy of the Upper

Carboniferous and Lower Permian in the

Donets Basin

Alain IZART Celine Briand "»», Denis VASLET 12 ', Daniel VACHARD' 3 ',

Jean BROUTIN 141 , Robert COQUEL 13 ', Alexander MASLO' 5 ',

Natalia Maslo 15 ' & Raissa Kozitskaya 151

Laboratoire de Geologie des ensembles sedimentaires, Universite de Nancy I

UMR CNRS 7566, B.P. 239, F-54506 Vandceuvre les Nancy, France

,21 BRGM - SGN/GEO, B.P. 6009. F-45060 Orleans Cedex 2, France

Laboratoire de Paleontologie, University Sciences et Technologies de Lille

URA CNRS 1365, F-59655 Villeneuve d'Ascq Cedex, France

l4) Laboratoire de Paleobotanique, Universite P. et M. Curie, 12, rue Cuvier, F-75005 Paris, France

l5> National Academy of Sciences of Ukraine, Institute of geological Sciences

55b Chaklov street, Kiev 252601, Ukraine

ABSTRACT

The Artemovsk geological survey published numerous logs for the Moscovian. Kasimovian, Gzhelian and Lower Permian

in a transitional area between continental and marine environments of the Donets Basin. New results concern the stratigraphy

according to fusulimds. plants and palynomorphs and sequence stratigraphy. They have been developed by comparison of our

observations on cross-sections near Artemovsk and the published coal mine logs. In the Donets Basin, the Kasimovian forms a

second order sequence as in the Moscow Basin, that can be subdivided into three third order sequences, which can be further

subdivided into fourth order sequences. The Kasimovian is equivalent to the lower Stephanian of the western European basins

The Gzhelian forms a second order sequence that can be subdivided into four third order sequences, which can be further

subdivided into fourth order sequences. The Gzhelian is equivalent to the upper Stephanian and the lowermost Autunian of the

western European basins. The Asselian is a second order sequence that shows a lowstand and a transgressive system tract

dunng the early and middle Asselian, a highstand system tract during the late Asselian, that continues during the Sakmarian. In

the Donets Basin, the fourth order sequences of the Carboniferous begin with an erosive base and show a succession of

elementary sequences with fluvial sandstone, coal-bearing, limestone and deltaic facies. The fluvial sandstone represents a

period of aggradation, the coal and the limestone a transgressive period and the deltaic facies a period of progradation. These

fourth order sequences have a higher amplitude than the third order sequences and obscure these latter sequences. The second

order sequence equivalent to the Kasimovian represents a lowstand system tract with fluvial sandstone and a transgressive

system tract with limestone in the early Kasimovian, a highstand system tract with deltaic facies in the middle and late

Kasimovian. The second order sequence equivalent to the Gzhelian presents a lowstand system tract with fluvial sandstone and

a transgressive system tract with limestone during the early and middle Gzhelian, a highstand system tract with deltaic facies

Izart, A.. Briand, C., Vaslet, D., Vachard. D., Broutin, J., Coquel, R., Maslo, A., Maslo, N. & Kozitskaya, R.,

1998.— Stratigraphy and sequence stratigraphy of the Upper Carboniferous and Lower Permian in the Donets Basin. In: S.

Crasquin-Soleau & E. Barrier (eds), Peri-Tcthys Memoir 3: stratigraphy and evolution of Peri-Tethyan platforms Mem

Mus. natn. Hist, nat.. Ill : 9-33. Paris ISBN : 2-85653-512-7.

Source: MNHN. Paris

ALAIN IZART ET AL.

during the late Gzhelian. The uppermost Gzhelian, equivalent to the lowermost Autunian, and the Asselian consist of red fluvial

deposits with some Autunian plants and lagoonal limestones. The sequences are controlled by regional tectonic subsidence in

the graben and the uplift of the horsts and / or by eustasy with either a plate-tectonic or a glacial origin. However, other

investigations, especially radiometric dating, will be useful for solving this problem. Each third order sequence corresponds to a

biozone of fusulinids. Lithological and biological sequences are identical. The thicknesses and tectonic subsidence decrease

gradually from the Moscovian to the Sakmarian. Hiatuses exist between the Sakmarian and Tatarian and between Tatarian and

Lower Triassic, corresponding to an uplift or a tectonic inversion period. The drift of the Pangea into equatorial latitude during

Carboniferous and tropical latitude near the Carboniferous / Permian boundary explains the differences in lithology (coal /

evaporites, red beds), plants (hygrophylous / xerophylous) between Carboniferous and Permian.

RESUME

Stratigraphic et stratigraphie sequentielle du Carbonifere superieur et du Permien inferieur dans le bassin du

Donetz.

Le service geologique d'Artemovsk a publie de nombreux logs pour le Moscovien, le Kasimovien, le Gzhelien et le Permien

inferieur dans une zone de transition entre le domaine continental et marin du bassin du Donetz. Les nouveaux resultats

concernent la stratigraphie d’apres les fusulines, la flore et la microflore. et la stratigraphie sequentielle. Ils ont ete obtenus en

comparant nos observations sur les coupes situees pres d’Artemovsk et les logs miniers. Dans le bassin du Donetz. les depots du

Kasimovien forment une sequence du deuxieme ordre comme dans le bassin de Moscou, qui peut etre subdivisee en trois

sequences du troisieme ordre. elles-memes subdivisees en sequences du quatrieme ordre. Le Kasimovien est equivalent au

Stephanien inferieur des bassins de I'Europe de l'Ouest. Les depots du Gzhelien forment une sequence du deuxieme ordre qui

peut etre subdivisee en quatre sequences du troisieme ordre. elles-memes subdivisees en sequences du quatrieme ordre. Le

Gzhelien est ('equivalent du Stephanien superieur et de la partie inferieure de 1'Autunien des bassins de I'Europe de l'Ouest. Les

depots de l'Asselien et du Sakmarien forment une sequence du deuxieme ordre qui presente le cortege de bas niveau et le

cortege transgressif pendant l’Asselien inferieur et moyen et le cortege de haut niveau pendant l'Asselien superieur et le

Sakmarien. Dans le bassin du Donetz, les sequences du quatrieme ordre du Carbonifere commencent avec une base erosive et

montrent une succession de sequences elementaires avec gres fluviatile, charbon, calcaire et facias deltai'que. Le gr£s fluviatile

represente une periode degradation, le charbon et le calcaire une periode transgressive et le facies deltai'que une periode de

progradation. Les sequences du quatrieme ordre ont une amplitude plus grande que les sequences du troisieme ordre et les

masquent. La sequence du deuxieme ordre du Kasimovien presente un cortege de bas niveau avec gr£s fluviatile et un cortege

transgressif avec calcaire pendant le Kasimovien inferieur, et un cortege de haut niveau pendant le Kasimovien moyen et

superieur. La sequence du deuxieme ordre du Gzhelien presente un cortege de bas niveau avec gres fluviatile et un cortege

transgressif avec calcaire pendant le Gzhelien inferieur et moyen, et un cortege de haut niveau avec des faci&s deltai'ques

pendant le Gzhelien superieur. La partie superieure du Gzhelien, encore appelee Orenburgien, equivalente de la partie inferieure

de 1'Autunien, et l'Asselien consistent en des depots fluviatiles et lagunaires rouges a plantes d'age Autunien. Les sequences

sont controlees par la subsidence tectonique regionale dans le graben du Donetz et le soulevement des horsts qui l'entourent et /

ou par 1'eustatisme soit avec une origine glaciaire soit liee a la tectonique des plaques. Cependant, d'autres investigations, en

particulier geochronologiques seraient utiles pour resoudre ce probleme. Chaque sequence du troisieme ordre correspond a une

biozone de fusulines. Les sequences lithologiques et biologiques sont identiques. Les epaisseurs et la subsidence tectonique

decroissent graduellement du Moscovien au Sakmarien. Des hiatus existent entre le Sakmarien et le Tatarien et entre le Tatarien

et le Trias inferieur, correspondant a un soulevement ou a une periode d’inversion tectonique. La derive de la Pangee au travers

des zones equatoriale pendant le Carbonifere et tropicale pendant le Permien expliquent les differences dans la lithologie

(charbon/evaporites, couches rouges), plantes (hygrophiles/xerophiles) entre le Carbonifere et le Permien.

INTRODUCTION

The Donets Basin, an Ukrainian coal-bearing zone, forms a part of the Pripyat-Dnieper-Donets rift

(Fig. la) that extends from the Baltic to the Caspian Sea across Belarus, Ukraine and Russia

(AlSENVERG et al ., 1975; Chekunov et al., 1993). This zone is situated between two basement horsts

that provide sediments: the Voronezh crystalline massif in the north and the Ukrainian crystalline massif

in the south. STEPHENSON et al. (1993), Stovba et al. (1995) discussed the rift development of the

Dnieper-Donets basin: pre- and syn-rift stages during the Devonian and Early Carboniferous, post-rift

stage with extensional rejuvenations during the Late Carboniferous, post-rift stage between the Permian

and Tertiary with an uplift in the Permian and a tectonic inversion in the Early Tertiary. During the

Carboniferous, the Donets basin was a paralic coal zone forming a transitional domain between the

paralic and limnic western European coal basins (e.g.: the North France paralic coal Basin and the Saint-

Etienne limnic coal Basin) and the sea on the Russian platform (IZART & VACHARD, 1994). Therefore,

the Donets basin is a key area for solving the problems of correlations between the continental and

marine deposits during the Carboniferous.

After our synthesis (IZART & VACHARD, 1994), this publication provides new data on the

sedimentology and stratigraphy in the Donets Basin which could be useful for improving global

Source: MNHN . Paris

STRATIGRAPHY OF UPPER CARBONIFEROUS AND LOWER PERMIAN IN DONETS BASIN

Fig. 1

Fig . I

. The Pripyat-Dnieper-Donets Basin. A: Sketch of the Prypiat-Dnieper-Donets Basin. 1. Pripyatsky graben- 2 Dnieper

crvstalIinp^rna« b k^* 4 ’- B £ rderS ? f J^asin ;‘ Voronezh crystalline massif; 6, Pripyatskaia anticline; 7, Ukrainian

crystalhne mass ,f ; 8, Karpinsky arch; 9. North Caspian syncline; 10, Sea. The rectangle represents the map B. B:

iqr?w. °v i hC cross_sect,ons . m the Donets Basin. AHI: Mine logs from Artemovsk geological survey (Makarov,

Redichkin 1980)° cross " secl,on slud,ed b y our team - KL: Kalmuys-Bakhmutskaya cross-section (Movskovich &

Le bassin Pripyat-Dnieper-Donetz, A : Carte geologique schematique du bassin Pripyat-Dnieper-Donetz, 1. graben

du Pripyat; 2, graben du Dnieper; 3, graben du Donetz; 4, bordures du bassin; 5, massif cristallin du Voronezh- 6

anticlinal de Pripyat; 7. massif cristallin d'Ukraine; 8, arche de Karpinsky; 9. synclinal Nord caspien; 10 Mer Le

rectangle represente la carte B. B Carte de situation des coupes du bassin du Donetz . AHI: logs miniers du service

geologique d Artemovsk ( Makarov ; 1985). J : coupe de Kalinovo etudiee par noire equipe. KL : coupe Kalmuys-

Bakhmutskaya (Movskovich & Redichkin, 1980). 1 '

Source:

ALAIN 1ZART ETAL.

correlations. In 1994. some logs of the Moscow platform have been surveyed with Russian

stratigraphers (BRIAND et al ., in press) and of the Donets Basin with Ukrainian stratigraphers of the

Kiev^Institute of Geological Sciences in a cross-section in Kalinovo (J. Fig.lb) for the Moscovian

(IZART et al., 1996), Kasimovian, Gzhelian and Lower Permian. Sedimentological and biostratigraphical

studies of the foraminifera, conodonts, plants and spores are now in progress. This publication deals

with the sequence stratigraphy and biostratigraphy of the Kasimovian. Gzhelian and Lower Permian of

the Donets Basin in a transitional zone between continental and marine areas and the geological controls

in this basin.

STRATIGRAPHY AND SEQUENCE STRATIGRAPHY OF THE KASIMOVIAN,

GZHELIAN AND LOWER PERMIAN OF THE DONETS BASIN

Stratigraphic outline

Ukrainian stratigraphers have published many contributions on the geology and stratigraphy of this

basin. ZHEMCHUZHNIKOV et al. (1959-1960) presented a cyclostratigraphy of the Moscovian of the

Donets Basin. AISENVERG et al. (1960, 1975) published on the stratigraphy of the Carboniferous of the

Donets. STSCHEGOLEV (1960, 1965, 1975) and BOYARINA (1994) published on Carboniferous and Early

Permian floras. MAKAROV (1985) presented many logs in the transitional area between continental and

marine deposits along the cross-section AHIJ for the Kasimovian-Gzhelian (Fig. lb). KOZITSKAYA et al.

(1978) published a biostratigraphic scale for conodonts of the Donets Basin. SOLOVIEVA (1985) and

SOLOVIEVA et al. (1985 a, b) proposed new stratigraphic correlations for the Donets Basin. DAVYDOV

(1990) published on Kasimovian and Gzhelian foraminifera. KAGARMANOV & DONAKOVA (1990)

proposed correlations between Moscow platform, Ural and Donets basins. MOVSKOVICH & REDICHKIN

(1980) presented a cross-section of the Lower Permian, while AISENVERG et al. (1975) and LAPTCHIK

(1970) published many logs of the Lower Permian in the Artemovsk area and Dnieper Basin.

AISENVERG et al. (1960, 1975) described lithostratigraphic units named suites (Table 1) in the Donets

Basin and these subdivisions have been chosen for the description. But, KAGARMANOV & DONAKOVA

(1990) named these suites horizons (Table 1) and change their boundaries. The horizons in the russian

terminology are identical with the Formations in the standard terminology (SALVADOR, 1994).

The table 1 shows the lithostratigraphy, the biostratigraphy and the chronostratigraphy of the Donets

Basin. Chronostratigraphic equivalences are proposed between the Ukraine, the eastern Europe (Russia)

and the western Europe.

Methodology

Facies data and paleoenvironments interpretations are based on the logs (e.g.: Fig. 2). For further

information refer to MAKAROV (1985) along the cross-section AHIJ (Fig. 1). The authors checked some

data in the field (J, Fig. 1) and further data will be presented by BRIAND (Thesis in progress).

AISENVERG et al. (1975) and MAKAROV (1985) distinguish in the Moscovian, Kasimovian and

Gzhelian of the Donets Basin: a- marine deposits (M) with marine limestones and claystones and marine

deltaic siltstones and claystones, b- transitional deposits with lagoonal or lacustrine claystones and

lagoonal or lacustrine deltaic siltstones or sandstones, c- swampy deposits with coal seams, d-

continental deposits (F) with fluvial sandstones. Marine bands are indicated by capital letters (e.g.: N,

Fig. 2) and coal seams with small letters (e.g.: n ). Deltaic deposits exhibit coarsening upward sequences

from shale to sandstone and fluvial deposits fining upward sequences from sandstone to shale. Channels

and cross-beddings occur in the fluvial sandstones.

A sequence stratigraphy has been established on the basis of lithological and paleoenvironmental

data and new field data. The environments were sorted out between two trends, one marine at the left

(M) and the other continental (F) at the right to exhibit the retrogradation and the transgression towards

the west and the progradation and the regression towards the east according to the location of the sea on

the Moscow platform.

Source: MNHN, Paris

STRATIGRAPHY OF UPPER CARBONIFEROUS AND LOWER PERMIAN IN DONETS BASIN

TABLE l ,V,- 7 ^ tr A t ; 8 i aphy ° f lhe K ;* si movian. Gzhelian and Lower Permian in the Donets Basin. (I) AlSENVERG el al. (1960

1975) (2) Kagarmanov & »)nakova (1990), (3) Vachard& Maslo (this note), (4) Donets local megafloral zones

according to Stschegolev & Kozitskaya (1984) and Boy arina (submitted), (5) this note.

PRL: Pecopteris arcuata , Raminervia mariopteroides, Lodevia nicklesii

ASA: Asterotheca daubreii, Sphenopteris gennanica, Alethopteris subelegans

ASE: Alethopteris subelegans , Sphenopteris gennanica, Emestiodendron fdiciformis

TABLEAU I — Stratigraphie du Kasimovien, Gzhelien et Pennien inferieur du bassin du Donetz . (I) AlSENVERG et al. (I960

I 9 75 ), (2) Kagarmanov & Donakova (1990), (3) Vachard& Maslo (ceiie note), (4) zones locales florales du Donetz

d apres Stschegolev & Kozitskaya (1984) et Boy arina (soumis), (5) cette note.

LITHOSTRATIGRAPHY

biostratigraphy

chronostratjgraphy

I Suites

( 1 )

Horizons

( 2 )

Fusulinids

(3)

Plants

(4)

Western F.urop*

Continental

"Stages" (5)

Eastern Europe

Marine Stages

(5)

Pj-kmi

(T)

Kramatorsky

"Sakmarella"

moelleri

Sakmarian

Pl-sl

(S)

Slavjansky

Sphaero-

schwagerina

sphaerica

As 3

Pj-nk

(R)

Nikitovsky

Sphaero-

schwagerina

moelleri

Aui uni a

Autunian

As 2

Lebachia

Pj-krt

(Q)

Kanamyshky

Sphaero-

schwagerina

fusiformis

As 1

ASE

Uliradaixira

Mironovsky

bosbyiauensis

Daixirui

ASA

c 3 p

03 <p)

robusta

PRL

Daixina

sokensis

c 3 e

Kalinovsky

Jiguliles

jigulensis

Asteroiheca

densifolia

Upper

Stephanian

c 3 d

C 3 (O)

Triiiciies

rorsicus

Triiiciies

siuckenhcrgi

Odoniopieris

osmundaeformis

C 3 C

Triiiciies

acuius

Triiiciies

quasiarciicus

c 3 b

Torezky

Monliparus

montiparus

Asterotheca

lamuriana

Lower

Stephanian

C 3 A 2

C 3 (N)

Proiriiicites

pseudo-

montiparus

Obsoleies

obsoleius

Neuropleris

oval a

-3 a 1

Praeobsoleies

burkemensis

Sanjarovsky

Fusulma

cylinlrica

Neuropleris

ovaia

Westphalian D

Moscovian

Marine

Bands

R 4

R I

Q11

P 5

0 4 5

0 |

N 5

n 4

n 3

Source

ALAIN IZART ETAL.

The authors use sequence stratigraphy (VAN WAGONER et al ., 1988) in a transitional area between a

more continental area westwards in the Dnieper Basin and a more marine area eastwards, because the

erosion surfaces are well exposed and the maximal flooding surfaces are well recognizable and placed at

the acme of the fusulinids biozone. In figures 2 and next, high-frequency sequences (about 40 Ka-100

Ka), fourth order (about 400 Ka), third order (about 1 Ma) and second order sequences are described. In

the sequence stratigraphy, a parasequence is defined between two flooding surfaces. But, the lirst

particular problem that was perceived by SHANLEY & McCabe (1994) for the application of sequence

stratigraphic concepts to continental strata was: "What, if anything, is the continental equivalent to a

parasequence?". In this type of paralic coal basin, the authors prefer to use the term elementary sequence

defined between two more continental trends, the base and the top of the sequence. An elementary

sequence is a high-frequency sequence, probably a fouth order sequence (about 100 Ka) or a fifth order

sequence (about 40 Ka). Each sequence presents either three facies / paleoenvironments: a- more

continental at the sequence base, b- more marine at the maximal flooding surface and c- more

continental at the sequence base of the upper sequence. Four facies types may also occur: a- plus b- plus

c- plus a coal transgressive facies. Locally, below the fluvial channel, an additionnal facies at the top ot

the delta shows the degree of erosion and the names of eroded beds are added between brackets. Various

cases can exist for the more continental trend: base of fluvial sandstone, seat earth at the base of coal

seam and for the more marine trend: limestone top or marine, lagoonal or lacustrine claystone.

The sequence continuity has been checked all along the cross-section by comparing the logs (Fig. 2).

This continuity is very reliable, because bore holes and underground mine exposures are numerous.

These were already synthesized by MAKAROV (1985). Especially marine bands are reliable stratigraphic

markers all across the Dnieper-Donets Basin. For the sake of simplicity, a stratigraphic correlation will

not be proposed for the elementary sequences. The marine bands can be followed all over the cross-

sections with a lateral passage between marine and lagoonal facies, except when there are eroded up¬

stream. Correlations are proposed for the fourth order sequences by junction-dash by using the high

degree of erosion at the base of channel for locating the base of the sequence. The order of each

sequence and the distinction between fourth, third and second order sequences are discussed in the text

according to the average duration of each sequence (CARTER et al., 1991, Table 2). From the detailed

figures 2 and next, it is difficult to observe third order and second order sequences, because a suite does

not correspond to a third order sequence. All the bore holes described by MAKAROV (1985) are used

here and figure 3 shows a simplified presentation of these sequences.

STRATIGRAPHY AND SEQUENCE STRATIGRAPHY OF THE Cj.N) SUITE

LlTHOSTRATIGRAPHY

Data have been recovered in the AH1J area (Fig. lb. Fig. 2) near Artemovsk (MAKAROV, 1985).

Three bore holes have been selected: 3924 in A, 4180 in H, 7045-7046 in I and the Kalinovo cross-

section in J. A cross-section in Kalinovo in J has been surveyed for checking the lithology and

environments. MAKAROV (1985) presented a cross-section composed of fifteen boreholes in AHIJ area.

This suite was defined between N, and O, (Table 1). The boundary between the Moscovian and

Kasimovian has been placed either in N, (AlSENVERG et al., 1960, 1975) or in N 4 (KAGARMANOV &

DONAKOVA, 1990 and SOLOVIEVA et al., 1985a).

Biostratigraphy and chronostratigraphy

Fusulinids.— This suite comprises two fusulinid biozones characterizing the upper Moscovian and

lower Kasimovian (Table 1). VILLA et al. (1993) reported the occurence of Protriticites in Nj . They

defined the Protriticites ovatus, Quasifusulinoides quasifusulinoides, Praeobsoleles tethydis biozone

between Nj and Nj which is an indeterminate interval attributable to the Moscovian or the Kasimovian.

The Moscovian Praeobsoletes burkemensis biozone defined by REMIZOVA (1995) could be represented

by the interval N 4 and Nj in the Donets Basin. VILLA et al. (1993) reported the occurence of Obsoletes

in Nj. The Protriticites pseudomontiparus and Obsoletes obsoletus biozone C,A, dated as Kasimovian is

correlatable with the interval between Nj and O] . Between N,and O, appears to evolve progressively

with a gradual transition from the Moscovian to the Kasimovian: the appearance of Quasifusulinoides in

Source:

STRATIGRAPHY OF UPPER CARBONIFEROUS AND LOWER PERMIAN IN DONETS BASIN

E|SW

SMI7

FACIES:

dog)

H Limestcne

| Gayslone

■”—Coil with palcosol

L-'V-A'-j Sandslonc

P AL EOE.VVIRONM ENTS:

(sequence diagram)

Marine

trend

1 11 Se»

m Marine della

| L | Lagoon-lake-della

1 S | Swamp

-i Conlincnlal

River trend

J m

Fig. 2.— Lithostratigraphy and sequence stratigraphy of the Cj (N) Suite in the southern transitional area AHI and in the

northern transitional area J in Kalinovo. The notations in capital letter represent the marine bands (e.g.: N,), in small

letter the coal seams (e.g.: n 0 ) and in brackets the name of eroded beds.

Fig. 2 — Lithostratigraphie et stratigraphie sequentielle de la Suite Cj (N) dans les zones de transition AHI et J a Kalinovo. Les

notations en lettres majuscules representent les niveaux rnarins (par exemple : N,), en lettres minuscules les veines de

charbon (par exemple : n 0 ) et entre crochets le nom des couches erodees.

Source: MNHN. Paris

ALAIN IZART ET AL.

N 2 , the disappearance of Fusulinella bocki and appearance of Fusulina cylindrica in N 3 , the appearance

of Protriticites and Praeobsoletes in N 4 , an acme of Protriticites in N 5 , the appearance of Tubiphytes,

Archaelithophyllum lamellosum and Obsoletes in NJ and an acme of Obsoletes in O,. The fusulinids

occuring between NJ and 0\ are the same as those in the Protriticites pseudomontiparus-Obsoletes

obsoletus C 3 A, biozone which is dated as early Kasimovian, i.e. the Krevyakinian of the Moscow Basin.

KOZITSKAYA et al. (1978) found a similar gradual transition in the conodonts and they defined the base

of the Kasimovian in NJ. As it is difficult to attribute the interval NJ - NJ to the Moscovian or the

Kasimovian, a sedimentological criterion, the highest erosion at the base of the sandstone between N 3 an

N 4 (see after the sequence stratigraphy), has been preferred to determine the limit between the

Moscovian and the Kasimovian.

Plants.— In Kartanash, STSCHEGOLEV in AlSENVER Get al. (1975, p. 232) and STSCHEGOLEV &

KOZITSKAYA (1984) found hygrophilous plants with " Asterotheca " lamuriana in n 3 below the

sandstone. And at the top of the sandstone, he collected xero- and mesophilous plants with Walchia (aL

Lebachia) piniformis , Culmitzschia (al. Lebachia) parvifolia, Cordaites sp. Silicified logs of Dadoxylon

sp. was found in the sandstones between 0 2 and 0 4 . If the Donets Basin is compared with the Saint-

Etienne Basin, only " Asterotheca lamuriana" is considered as early Stephanian. Most of other plants are

rather long-ranging and not restricted to the Stephanian. As the plants are absent between N 3 and N 4 , the

same sedimentological criterion as for the Moscovian has been used to determine the limit between the

Westphalian and the lower Stephanian. The lower Stephanian extends thus from the base of the

sandstone between N 3 and N 4 to 0 5 and is equivalent to the Kasimovian according to these new

identifications.

Sequence stratigraphy (Figs 2, 3)

High-frequency and fourth order sequences. — Six sequences are proposed for this suite

(Figs 2, 3): three attributed to the upper part of the Myachkovian (SMI6 to SMI8) and three to the lower

Kasimovian (SKI to SK3). The duration of the Kasimovian was about five million years, the duration of

these individual sequences is less than one million years and they are therefore fourth order sequences.

The sequences SM6 to SMI8 present elementary sequences with four facies / paleoenvironments: a-

sandstone (fluvial) - b- coal (swampy) - c- limestone (marine) - d- coal (swampy) during the aggradation

and the retrogradation, or three: a- coal (swampy) - b- claystone (lagoonal) - c- coal (swampy) during

the progradation. The sequences SKI to SK3 show elementary sequences with four facies /

paleoenvironments: a- sandstone (fluvial) - b- coal (swampy) - c- limestone (marine) - d- sandstone

(fluvial) during the aggradation, the retrogradation and the progradation. So, these elementary sequences

gather in a stacking pattern with erosion at the base, an aggradation for the fluvial sandstones in a

lowstand system tract. The erosion is low at the base of SMI6 and SMI7, strong at the base of SMI8

(Fig. 3), where Nj 7 and N* are eroded in the bore holes 3924, 4180, 7045. The erosion is strong at the

base of SKI as NJ is eroded in the bore holes 3924, 4180, 7045, strong at the base of SK2 as NJ, N 3 3 are

partly eroded in 3924, strong at the base of SK3, as NJ is partly eroded in 3924 and the top of the deltaic

deposits in the other bore holes. A retrogradation in a transgressive system tract is observed between the

flooding surface and the maximal flooding surface, and a progradation in a highstand system tract

between the maximal flooding surface and the base of the next sequence. The flooding surface is

positioned at the first coal seam of the sequence in no for SMI6, in nj for SMI7, in n] for SMI8, in nj

for SKI, in nj below N 4 or SK2 and in nj for SK3. The maximal flooding surface is placed at the top of

the more transgressive limestone of the sequence with some uncertainty in comparing the logs, in N, 1 to

NJ for SMI6, in N? for SMI7, in N 2 for SMI8, in NJ b for SKI, in N 4 for SK2, in O, for SK3. The more

transgressive limestone is the limestone that presents the highest spatial amplitude in the basin (see

figure 3). For each sequence, these three system tracts thus form a fourth order sequence.

The thickness of these sequences increases towards the east and decreases towards the north, the

thickness of sandstone increases westwards and the marine deposits increase eastwards. An evolution of

the degree of erosion can be observed at the base of the Kasimovian: middle below N 2 , high below NJ b

and N 4 . In conclusion, the limit would be between N 3 and N 4 .

Third ORDER SEQUENCE. — The fourth order sequences in the upper part of this suite form a third

order sequence. A fluvial lowstand system tract is positioned between the erosive surface on N 3 and the

coal seam nj. A transgressive system tract is placed between the flooding surface nj and the maximal

Source: MNHN. Paris

STRATIGRAPHY OF UPPER CARBONIFEROUS AND LOWER PERMIAN IN DONETS BASIN

flooding surface O,. A highstand system tract is positioned between O, and the base of the next

sequence. O, was chosen according to the appearance of the first typical Kasimovian foraminifera in this

marine band. This sequence is dated as early Kasimovian or Krevyakinian in Moscow Basin.

STRATIGRAPHY AND SEQUENCE STRATIGRAPHY OF THE C 3 2 (0> SUITE

LITHOSTRA T/GRA PHY

Data are recorded from the AHIJ area (Fig. lb, Fig. 4) near Artemovsk (MAKAROV, 1985). Three

bore holes have been selected: 3881 in A, A441 in H, 7048 in I and the Kalinovo cross-section in J. A

cross-section in Kalinovo in J has been surveyed for checking the lithology and environments.

Makarov (1985) presented a cross-section composed of fifteen boreholes in the AHIJ area. This suite

was defined between O, and F, by Ukrainian stratigraphers. And the boundary between the Kasimovian

and Gzhelian has been placed variously in P, (AlSENVERG etal, 1960, 1975) or in 0 6 (KAGARMANOV &

Donakova, 1990 and Solovieva et al, 1985a).

Biostratigraphy and chronostrat/graphy

FUSULINIDS.— This suite corresponds to three Fusulinid biozones characterizing middle

Kasimovian, upper Kasimovian and lower Gzhelian (Table 1). VILLA et al (1993) reported the

appearance of Montiparus in Oj , a limestone under 0 2 . They defined the Montiparus montiparus

biozone C 3 A 2 between Oj and 0 4 . Montiparus and Quasifusulina longissima occur in Q>; there is an

acme of Montiparus in 0 3 . These fusulinids correspond to the biozone C 3 A 2 and indicate a middle

Kasimovian age, the Khamovnichian in the Moscow Basin.

Davydov (1990) and Villa et al (1993) observed Triticites quasiarcticus in 04, 0 4 ' and SEMINA

(1984) found Triticites acutus in 0$. The first Triticites without specific identification occur in Oj and

OJ. These fusulinids correspond to the Triticites acutus-Triticites quasiarcticus biozone B and indicate a

late Kasimovian age, the Dorogomilovian and Yausian in the Moscow Basin.

Davydov (1990) reported the occurence of Nodosaria cf. ronda and Triticites rossicus in 0 5 and 0 6 .

Our identifications confirm these data. These foraminifera correspond to the Triticites rossicus-Triticites

stuckenbergi biozone QC and indicate a early Gzhelian age, the Rechitskian and Amerevian in the

Moscow Basin. Davydov (1990) reported here also the first Jigulites and Daixina corresponding to the

Jigulites altus subzone. A sedimentological criterion, the highest erosion at the base of the sandstone

between OJ and OJ (see after the sequences stratigraphy) has been preferred to determine the limit

between the Kasimovian and the Gzhelian.

PLANTS. — This suite corresponds to two plant biozones characterizing lower Stephanian and upper

Stephanian. STSCHEGOLEV i n AlSENVERG etal. (1975) and STSCHEGOLEV & KOZITSKAYA (1984)

found " Asterotheca" lamuriana and Neuropteris ovata up to 0 5 . Neuropteris ovata ranges from the

Westphalian D to early Stephanian in the North France Basin (Laveine, 1986) and in the Saint-Etienne

Basin (DOUBINGER et al , 1995, table 8). In the Donets Basin, the lower Stephanian continues thus to O s

and is equivalent to the Kasimovian according to our new identifications.

STSCHEGOLEV, in AlSENVERG et al ., 1975) found between 0 5 and 0 7 hygrophilous plants with

Asterotheca densifolia. This taxon is known in the upper Stephanian of the Saint-Etienne Basin

(DOUBINGER et al , 1995). The same sedimentological criterion as for the Gzhelian has been used to

determine the base of the upper Stephanian placed at the base of the sandstone between OJ and OJ . The

upper Stephanian extends thus from this sandstone to p 4 and is equivalent to the Gzhelian pro parte.

Sequence stratigraphy

HIGH-FREQUENCY and FOURTH ORDER SEQUENCES. — Seven sequences are proposed here for this

suite (Fig. 4, Fig. 3) with one attributed to the middle Kasimovian (SK4), three to the upper Kasimovian

(SK5 to SK7) and three to the lower Gzhelian (SGI to SG3). The duration of the Kasimovian was about

five million years, and of the Gzhelian, about five million years, the duration of these individual

sequences is less than one million years, they are therefore fourth order sequences.

W IJ

ALAIN IZART ETAL.

4 Ih ORDER

3rd ORDER

2ndORDER

SGI3

SGIO.

SMI 8

MOSCOVIAN

SM16

Sequence boundary

HST Highstand system tract

TST Transgressive system tract

sandstones

Marine bands

Flooding surface

Maximum flooding surface LST Lowstand system tract

Cj (P)

wsw

CjlN)

SOm

In km

SG3

HST

SG2

TST

~ ~ sgT ~

LST

SK6

HST

SK5

■■•■TSF- -

SK4

HST

— . — —

i j

-LL

TST

SK2

SKI

LST

HST

' Sequence stratigraphy of the Kasimovian and Gzhelian in the Donets Basin

t!C. 3.- Stratigraph,e sequentielle du Kasimovien et du Gzhelien du bassin du Doneti

Source: MNHN, Paris

~ <C & z > — r w x n o

STRATIGRAPHY OF UPPER CARBONIFEROUS AND LOWER PERMIAN IN DONETS BASIN

The sequences SK4, SK7, SGI and SG2 present elementary sequences with four facies /

paleoenvironments: a- sandstone (fluvial)- b- coal (swampy) - c- limestone (marine) - d- coal (swampy)

during the aggradation, the retrogradation and the progradation. The sequences SK5, SK6 and SGI

present elementary sequences with four facies / paleoenvironments: a- sandstone (fluvial) - b- coal

(swampy) - c- limestone (marine) - d- coal (swampy) during the aggradation and the retrogradation or

three: a- coal (swampy) - b- claystone (lagoonal) - c- coal (swampy) during the progradation. So, these

elementary sequences gather in a stacking pattern with erosion at the base, an aggradation for the fluvial

sandstones in a lowstand system tract. The erosion is low at the base of the sequences SK4, SK5, SK7,

strong at the base of the sequence SK6, as O? is eroded in bore holes 3881 and A441. The erosion is

strong at the base of the sequence SGI, SG2 as Of and 0 5 are eroded in the bore hole A441, and in SG3

as is eroded in bore holes A441 and 3881. A retrogradation in a transgressive system tract is observed

between the flooding surface and the maximal flooding surface; a progradation in a highstand system

tract between the maximal flooding surface and the base of the next sequence. The flooding surface is

placed at the top of the sandstone in the sequences SK7 and SGI and of the first coal seam of the

sequence in o, for SK4, in of for SK5, in o( for SK6, in o 5 , for SG2 and o] for SG3. The maximal

flooding surface is positioned at the top of the more transgressive limestone of the sequence with some

uncertainty in comparing the logs, in O, for SK4, in 0 4 for SK5, in Of for SK6, in Of for SK7, in Of for

SGI, in 0 6 for SG2, 0 7 for SG3. For each sequence, these three system tracts thus form a fourth order

sequence.

The thickness of these sequences increases towards the east and decreases towards the north. The

thickness of sandstone increases westwards and northwards and the marine deposits increase eastwards.

A strong erosion can be observed at the base of the Gzhelian (SGI to SG3). In conclusion, the Gzhelian

limit would be below O s .

Third ORDER SEQUENCES. — The fourth order sequences in this suite form three third order

sequences (Fig. 3). The first begins below O, and consists of one fourth order sequence dated as middle

Kasimovian by toraminifera. Other fourth order sequences have been eroded by the erosive surface

below 0 4 . A fluvial lowstand system tract is placed between the erosive surface on O, and the coal seam

o,. A transgressive system tract is positioned between the flooding surface o, and the maximal flooding

surface 0 2 . A highstand system tract is placed between the maximal flooding surface O, and the base of

the next sequence below 0 4 .

The second third order sequence begins below 0 4 and consists of three fourth order sequences dated

as late Kasimovian by foraminifera. The lowstand system tract is placed between the erosive surface

below 0 4 and the top of the sandstone, the transgressive system tract between the top of the sandstone

and the maximal flooding surface 0 4 . the highstand system tract between 0 4 and the base of the next

sequence below 0 4 .

The third third order sequence consists of three fourth order sequences dated as early Gzhelian. A

fluvial lowstand system tract is placed between the erosive surface on OJ and the top of the sandstone.

A transgressive system tract is positioned between the flooding surface at the top of the sandstone and

the maximal flooding surface 0 7 . A highstand system tract is placed between the maximal flooding

surface and the base of the next sequence.

Stratigraphy and sequence stratigraphy of the C 3 V> and the lower

PART OF THE P,.KRT SUITE

LlTHOSTRA TIGRA PH Y

Data are recorded from the AHIJ area (Fig. lb. Fig. 5) near Artemovsk (Makarov, 1985). Three

bore holes have been selected: A4756 in A, A4 in H, A530 in I and the Kalinovo cross-section in J. A

cross-section has been surveyed in Kalinovo in J for checking the lithology and environments.

MAKAROV (1985) presented a cross-section composed of fifteen boreholes in AHIJ area. This suite was

defined between P, and Q, by Ukrainian stratigraphers and the boundary between Gzhelian and Asselian

was placed in Q, by AlSENVERG et al. (1960, 1975). KAGARMANOV & DONAKOVA (1990) and

Solovieva et al. (1985a). However, new data published by DAVYDOV (1990, 1992) favour the position

of the the Gzhelian / Asselian boundary in Q s . The lower part of Pj.krt, the Kartamyshky suite will be

ALAIN IZART ETAL.

FlG ' and , SCqu ® n , ce strat j8 ra phy of the q,o, Suite in the southern transitional area AHI and in the

„ . northern transitional area J in Kalinovo. Legends and comments as in figure 2

lt!!2'! a ',T ap Ue e! straU & ra P hi f "Ventielle de la Salle Cj,o> dans les zones de transition AHI el J a Kalinovo Us

legendes el les commentaires sont les mimes que pour la figure 2.

Source: MNHN. Paris

STRATIGRAPHY OF UPPER CARBONIFEROUS AND LOWER PERMIAN IN DONETS BASIN

described here too. Data on the P, suites near Artemovsk (profile, Fig. lb) are recorded according to

AlSENVERG et al (1975, Fig. 65), MOVSKOVICH & REDICHKIN (1980, Fig. 2, 1984), NESTERENKO &

KIREEVA (1980) and Laptchik (1970). The Ivano-darievka cross-section has been surveyed,

comparison with the Pre-Donets Basin in bore hole 4199 (Fig. la) has been attempted by Davydov

(1990) and ALEKSEEVA et al. (1983). Figures 6 and 7 present the lithostratigraphy and the sequence

stratigraphy of the P, Suites.

BIOSTRATIGRAPHY AND CHRONOSTRAT/GRAPHY

FUSULINIDS. — These suites correspond to three Fusulinid biozones characterizing the middle and

upper Gzhelian (Table 1). ALEKSEEVA et al. (1983) reported Daixina ruzhencevi from P, and DAVYDOV

(1990) observed Schagonella minor in P 2 . Numerous specimens of Schagonella and few Jigulites were

found; however Jigulites jigulensis appears to be absent. These fusulinids are the same as those found in

the Jigulites jigulensis biozone C 3 D dated as middle Gzhelian, the Pavlovo-Posadian in the Moscow

Basin.

Davydov (1990) observed Daixina spp. in P 3 , Daixina sokensis was found in P 3 and ALEKSEEVA et

al. (1983) reported Daixina enormis and Daixina spp. from P 4 . These fusulinids are typical for the

Daixina sokensis biozone C 3 E, dated as late Gzhelian, the Noginskian in the Moscow Basin.

No fusulinids were found in the lagoonal limestones near Artemovsk. However, Alekseeva et al.

(1983) and Davydov (1990, 1992, 1993) studied the fusulinids of the bore hole 4199 (Fig. 6), located

in the more marine Pre-Donets area eastwards (Fig. 1). They proposed correlations between the two

areas: Q. is the same, but Q, in Pre-Donets is equivalent to Q in Artemovsk. The Gzhelian / Asselian

boundary is the subject of numerous discussions. Davydov (1990) put the Ultradaixina bosbytauensis -

Daixina robusta biozone C 3 F at the Asselian base. In 1992, he dated this biozone as late Gzhelian, the

Melekhovian in the Moscow Basin and the first Asselian biozone is therefore the Sphaeroschwagerina

fusiformis biozone Asl. The boundary between these two biozones represents the Gzhelian / Asselian

boundary (DAVYDOV, 1992) and the potential Carboniferous / Permian boundary (JlN Yu-Gan et al .,

1994). On the other hand, paleomagnetic data show a short-term normal polarity subzone (KHRAMOV,

1963) in the top of C 3 F in Donets and Ural (KHRAMOV & DAVYDOV, 1984, 1993). In the Donets Basin,

the normal polarity subzone is in the level of Q. / Q, and corresponds to the C 3 F top according to

SCHNEIDER et al. (1995). Note that there are uncertainties for the upper limit of this biozone in the

Donets Basin. According to Alekseeva et al. (1983), Q 9 in bore hole 4199 contains

Sphaeroschwagerina cf. fusiformis. Davydov (1992) placed the base of C 3 F in P 5 and the top in Q 9 in

bore hole 4199, being equivalent of Q 7 in Artemovsk. However, SCHNEIDER et al. (1995) regarded Q, as

the base of C 3 F and Q 5 as its top. We prefer an intermediate position with the base in P 5 and the top in

Q 5 , because this proposal takes into consideration the fusulinids with evolutive arguments and the

paleomagnetism.

PLANTS.— For the plants, these suites correspond to three plants biozones characterizing upper

Stephanian and Autunian. BOYARINA (submitted) defined below P^ a megafloral local zone AO with

Asterotheca densifolia and Odontopteris osmundaefonnis. Asterotheca densifolia characterizes the upper

Stephanian in the Saint-Etienne Basin (DOUBINGER et al ., 1995) and the Donets Basin. For the

palynology, INOSOVA et al. (1976) defined IX and X palynological zones between P 2 and P° 5 with

Stephanian spores and 5% Potonieisporites sp. dated as late Stephanian-Autunian. The coal p 2 below P 3

shows a dominance of "Stephanian" forms with some Potonieisporites sp. A dominance of "Autunian"

forms with Potonieisporites sp.is found in the claystone below the coal p 3 . Furthermore, a claystone

within the coal p 3 below P 4 shows a dominance of "Stephanian" palynomorphs with Schopfipollenites

ellipiticus , Columinisporites sp. and Potonieisporites sp. The age is regarded to be late Stephanian with

a coexistence of "Stephanian" forms in the coal and "Autunian" palynomorphs in the flood plain

claystone. This coexistence of "Stephanian" and "Autunian" forms previously described from western

Europe (BROUTIN et al , 1986; BROUTIN et a /., 1990; BECQ-GlRAUDON, 1993) is now also recorded for

the upper Stephanian of the Donets Basin. This coexistence is controlled by paleoenvironmental

variations. The "Stephanian" flora consists of hygrophilous elements growing in marshes within the

basin. The "Autunian" flora is typified by meso-xerophilous elements growing on topographic highs

bordering the basin. The "Autunian" pollen came mainly by air and accumulated in the flood plain, lake

or lagoon within the basin.

ALAIN IZART ETAL.

T 5 ' ““ “ "» *”*“» ™ ■»■> in the northern trnn.itiona,

*t£5ES

Source: MNHN. Pahs

STRATIGRAPHY OF UPPER CARBONIFEROUS AND LOWER PERMIAN IN DONETS BASIN

STSCHEGOLEV (1960, 1965, 1975), STSCHEGOLEV & KOSITSKAYA (1984) and Boyarina (1994,

submitted) found in the claystone below P 5 (Kalinovo, Fig. 5) meso-xerophilous Callipterid

Pteridosperms. Abundant Raminervia mariopteroides grew in a moderately humid flood plain. Scarce

Rhachiphyllum schenkii , Dichophyllum cuneata , Autunia conferta , Lodevia nicklesii grew in an elevated

less humid part of the flood plain. In the roof of p 5 , they found hygrophilous plants with Asterotheca

arborescens and Annularia stellata. Boyarina (submitted) defined megafloral local zones PRL

between P° and P 6 and ASA between P ft and Q 4 PRL is composed of Pecopteris arcuata , Raminervia

mariopteroides and Lodevia nicklesii. ASA is composed of Asterotheca daubreii , Sphenopteris

germanica and Alethopteris subelegans. Raminervia mariopteroides and Dichophyllum cuneata are taxa

not known from other basins. Autunia conferta and Lodevia nicklesii are also known in some layers of

the upper Stephanian in the Saint-Etienne Basin (DOUBINGER et al. , 1995). Therefore, this list of plants

is too local and an age can not be attributed to these plants that can be upper Stephanian or Autunian.

Based on the palynology, INOSOVA et al (1976) defined the zone XI between P° and Q, with 10%

Potonieisporites sp. and the zones I to IV of the Permian between Q, and Q s with 10% Punctatisporites

confusus. The claystone below p 5 , where Stschegolev described Callipterids, shows a dominance of

"Autunian" forms with Potonieisporites spp. (above 80 %) and striate bisaccate pollen grains

(Protohaploxypinus spp.). The bituminous coal p 5 , the last coal of the Donets Basin, is characterized by

a dominance of three "Stephanian" forms: Calamospora sp., Laevigatosporites sp., Endosporites

globiformis. Furthermore, the claystone below P 7 shows "Stephanian" forms with Crassispora kosankei ,

Tripartites aductus, Lundbladispora sp. and pollen grains, e.g. Marsupipollenites triradiatus ,

Gardenasporites sp., striate bisaccates are absent. These beds are dated as Autunian with a coexistence

of "Stephanian" and "Autunian" taxa, associated with the same paleoenvironmental variations during the

late Stephanian and Autunian of the western Europe (BROUTIN et al. , 1990). INOSOVA et al (in

AlSENVER Get al. , 1975, Fig. 6) and INOSOVA et al. (1976) showed the dominance of Autunian forms in

Q 8 , that is positioned near the Asselian base. The climatic change recorded at the limit between the

Carboniferous and the Permian from wet to dry conditions with xerophilous taxa already occuring in the

uppermost Carboniferous is evident in the Donets Basin like in the western Europe (BROUTIN et al .,

1990). So, there are two possibilities for the limit between Stephanian and Autunian: the first appearance

of xerophilous plants in p 5 or their acme in Q 8 . We choose the first solution. Such a choice implies that

Lowermost Autunian would be latest Carboniferous in age. In conclusion, the base of the Autunian and

of the C 3 F zone of the upper Gzhelian are identical. And the base of Autunian and of the normal polarity

subzone are identical. This solution agrees with MENNING et al. (1988) and SCHNEIDER et al. (1995)

that reported the normal polarity subzone inside the lower Rotliegend near the base of the Manebach

Formation in the Saale Basin, equivalent to the top of the Kusel Formation in the Saar Nahe Basin and

lower part of the Autunian in the Autun Basin.

Sequence stratigraphy

High-frequency and fourth order sequences.— Nine sequences are proposed here for the CJ

suite (SG4 to SGI2, Fig. 5, Fig. 3) and four (SGI3 to SGI6, Fig. 7) for the P,.krt suite attributed to the

Gzhelian. The duration of the Gzhelian was about five million years, the duration of these individual

sequences is less than one million years, they are therefore fourth order sequences.

The sequences SG4 to SGI3 present elementary sequences with four facies / paleoenvironments: a-

sandstone (fluvial) - b- coal (swampy) - c- limestone (marine) - d- coal (swampy) during the aggradation

and the retrogradation or three: a- coal (swampy) -b- claystone (lagoonal) - c- coal (swampy) during the

progradation. The sequences SGI4 to SGI6 are composed of one or several elementary sequences with

three facies / paleoenvironments: a- claystone (fluvial plain) - b- limestone (lagoonal) - c- claystone

(fluvial plain) during the aggradation, the retrogradation and the progradation. Between Q, and Q 6 , the

marine trend is located at the top of the limestone and the continental trend at the base of the sandstone

or in the middle of the claystone, because sandstone exists in this place in the Dnieper Basin (LAPTCHIK,

1970). So, these elementary sequences gather in a stacking pattern with erosion at their base, an

aggradation for the fluvial sandstones in a lowstand system tract. There is little erosion at the base of the

sequences SG5 and SG6 and of the sequences SG10 to SGI 6, strong erosion at the base of the sequences

SG7, SG8 and SG9, as P 5 is eroded in the bore holes A530, A4, A4756. They include red fluvial

sandstones between P 4 and Q,, red claystones between Q 2 and Q, in Artemovsk area and fluvial

sandstones between P 4 and Q 6 in the bore hole 4199. Then, a retrogradation in a transgressive system

ALAIN IZART ETAL.

SW DONETS (BAKMUTSKAYA) PRE-DONETS NE

Kramatorsko- Artemovskoe Novokarfagenskoe Ivanodarievka Bore hole

Chasovyarskaya (Sverdlov mine)

Fig. 6.— Lithostratigraphy and sequence stratigraphy of the upper Gzhelian and Lower Permian in the Donets Basin and Pre-

Donets Basin in the bore hole 4199. M: marine environment. L: lagoonal environment. C: continental environment.

Fig. 6 .— Lilhostratigraphie et stratigraphie sequentielle du Gzhelien superieur el du Permien inferieur dans les bassins du

Dunetz et du Pre-Donetz (forage 4199). M: environnement marin. L : environnement lagunaire. C : environnement

continental.

Source: MNHN, Paris

STRATIGRAPHY OF UPPER CARBONIFEROUS AND LOWER PERMIAN IN DONETS BASIN

tract is observed between the flooding surface, and the maximal flooding surface and a progradation in a

highstand system tract between the maximal flooding surface and the base of the next sequence. The

flooding surface is placed at the top of the sandstone for the sequence SG4 and on the first coal seam of

the sequence in p| for the SG5, in p 2 for SG6, in p, for SG7, in the top of the sandstone for SG8, in p 4 for

SG9, in the top of the sandstone for SG10 to SGI2 and in the base of the limestone for SGI3 to SGI6.

The maximal flooding surface is positioned at the top of the more transgressive limestone of the

sequence with some uncertainty in comparing the logs, in F, for SG4, in P 2 for SG5, in P, for SG6, in P 4

for SG7, in P 5 for SG8, in PJ for SG9, in P‘ for SG10, in P^ for SGI 1, in P 7 for SGI2, in Q, for SGI3, in

Q 3 for SGI4, in Q 4 for SGI5 and in Q 5 for SGI6. So, for each sequence, these three system tracts form a

fourth order sequence.

The thickness of these sequences increases towards the east and decreases towards the north, the

thickness of sandstone increases westwards and northwards and the marine deposits increase eastwards.

A strong erosion can be observed at the base of the fourth order sequence SG9, corresponding to the

Autunian base.

THIRD ORDER SEQUENCES.— The fourth order sequences in these suites form three third order

sequences (Fig. 3) belonging to the Gzhelian. The first is composed of two fourth order sequences. A

fluvial lowstand system tract is placed between the erosive surface on and the top of the sandstone. A

transgressive system tract is positioned between the flooding surface and the maximal flooding surface

P,. A highstand system tract is placed between the maximal flooding surface and the base of the next

sequence.

The second third order sequence is composed of two fourth order sequences. A lowstand system tract

is placed between the erosive on P 2 and top of the sandstone. A transgressive system tract is positioned

between the flooding surface and the maximal flooding surface P 3 . A highstand system tract is placed

between the maximal flooding surface and the base of the next sequence below P 5 .

The third third order sequence presents a strong erosion on P 5 , but the lowstand system tract is

positioned between the erosive surface on P 4 and the coal seam p 4 according to the paleontological

argument. A transgressive system tract is placed between the flooding surface p 4 and the maximal

flooding surface Q,. A highstand system tract is positioned between Q, and the base of the next sequence

in Q 6 . The extension of the marine bands (Fig. 7) is lower between P 5 and Q„ maximum in Q, and

decreases between Q, and Q s .

STRATIGRAPHY AND SEQUENCE STRATIGRAPHY OF THE LOWER PERMIAN SUITES

LlTHOSTRA TIGRAPHY

The same data as for P,.krt are used here (Figs 6-7). Ukrainian stratigraphers defined the P,.krt suite,

named Kartamyshky, between Q, and R,, the P,.nk, named Nikitovsky, between R, and S„ the P,.sl,

named Slaviansky, between S, and T„ the P.-krm, named Kramatorsky. P.-krm, P,.nk and P,.sl are

Asselian and P,.krm is lower Sakmarian. The Dronovsky suite, composed of speckled claystone and

sandstone is Tatarian (Upper Permian) and rests with an angular unconformity on the Sakmarian. Then,

the conglomerate of Lower Triassic rests with an unconformity on the Tatarian (LAPTCFIIK, 1970). So, a

hiatus corresponding to an uplift or a compressive phase is located between the Sakmarian and Tatarian

and another one between the Tatarian and Lower Triassic.

Biostratigraphy and chronostrat/graphy

FUSULINIDS.— LAPTCHIK (1970) described the Foraminifera of the Donets Basin. ALEKSEEVA et ai

(1983) and DAVYDOV (1990, 1992) recognized the Sphaeroschwagerina fusiformis biozone Asj (Table

1) between Q 7 and R, (lower Asselian in the Moscow and Ural basins). Davydov found also the

Sphaeroschwagerina moelleri biozone As 2 between R 2 and S, (middle Asselian), the

Sphaeroschwagerina sphaerica biozone As, from S, (upper Asselian). However, the Sakmarella

moelleri biozone (lower Sakmarian) from T, is more problematical. In the Donets Basin near

Artemovsk, fusulinids are absent because marine bands are dolomitic, except lor Q,, where

Sphaeroschwagerina constans of the As 2 biozone was found.

ALAIN IZART ETAL.

Plants. _ The plants are rare in these red beds and indicate a xerophilous flora. BOYARINA

(submitted) defined local floristic zones ASE between Q 4 and Q s and AL between Q s and R 4 . The zone

ASE is typified by Alethopteris subelegans , Sphenopteris germanica , Emestiodendron filiciformis. The

zone AL includes Autunia conferta and Culmitzschia (al. Lebachia) angustifolia In the Autun Basin,

BROUTIN et cd. (1986) observed also the abundance of the xerophylous plants with Autunia conferta as

in the Donets basin, the age of these taxa is Autunian unreservedly. INOSOVA et al. (1976) defined zones

V and VI between Q 4 and Q s and zones VII to IX between Q s and R,,. They noted more 10%

Potonieisporites novicus and 5% Vittatina in zones V and VI. They showed a dominance of Autunian

forms in Q 8 , placed near the Asselian base with Vittatina (up to 10%) and Potonieisporites novicus (10

to 50%). The climatic change in the Lower Permian results in the depositon of evaporites and red beds

with a xerophilous flora.

Sequence stratigraphy

HIGH-FREQUENCY AND FOURTH ORDER SEQUENCES.— Eleven sequences (SAi to SA i 1, Figs 6-7)

are proposed here for the Asselian and two for the Sakmarian (SSL SS2). The duration of the Asselian

was about six million years (MENNING, 1995), the duration of these individual sequences is less than the

million years, they are therefore fourth order sequences.

In the Bakhmutskaya area of the Donets Basin, near Artemovsk, the sequences SAI to SAI 1 begin in

the middle of the claystone, passing laterally in a sandstone in the Dnieper Basin. For instance, the SAI

base corresponds to the palynological biozone V base below Q„ (Fig. 7). In the bore hole 4199 of the

Pre-Donets, the sequence SAI begins at the base of the sandstone below Q 5 , SA2 below Q l0 and

sequences can not be proposed for the limestones above without precise facies descriptions.

In the Donets Basin, these sequences represent one elementary sequence or several (e.i. Q 7 -Q 9 )- They

show a lowstand system tract and a transgressive system tract with red claystone in Q and R marine

bands, green claystone in S marine bands or halite and potash, a maximal flooding with gypsum,

limestone or dolomite, a highstand system tract with claystone. In the Pre-Donets Basin, the deposits are

only limestones as in the Moscow and Ural basins. In the Dnieper Basin. LAPTCHIK (1970) reported

sandstone at the Q,. R„ S, and T, base, sandstone in P,.krt, gypsum, salt, dolomite and limestone in Pi¬

nk, P,.sl and P|-krm.

Third ORDER Sequences.— The fourth order sequences can be gathered in three third order

sequences for the Asselian corresponding to the biozones As,, As, and As 3 and one for the Sakmarian.

The first third order sequence presents a lowstand system tract below Q 6 with a minimum extension of

the marine bands (Fig. 7). a transgressive system tract between Q 6 and Q„ a maximal Hooding surface in

Q, corresponding to the maximum extension, a highstand system tract in Q s , Q„, Q, 0 with minimum

extension. Note that for the sequence stratigraphy, there are two possibilities for the Asselian base

corresponding to a minimum extension in Q< or in Q„. The second third order sequence presents a

lowstand system tract in Q u -Q l2 with a minimum extension of the marine bands, a transgressive system

tract between Q,, and R,, a maximal flooding surface in R, known in the Dnieper Basin and

corresponding to the maximum extension, a highstand system tract in R,, R 3 , R 4 with minimum

extension. The third third order sequence does not present a lowstand system tract and a transgressive

system tract, but presents a maximal flooding surface in S, corresponding to the maximum extension and

a highstand system tract between S, and T, with minimum extension.

The Sakmarian comprises only a single third order sequence that continues the highstand system tract

of the Asselian before the closing of the Donets basin by uplift or compression, when eustatism

continues in the Ural.

Discussion on third and second order sequences

Figures 3 and 7 present the sequence stratigraphy of the Kasimovian, Gzhelian and Asselian of the

Donets Basin. The third order sequences of the Upper Carboniferous are arranged in two second order

sequences that are equivalent to the Kasimovian and Gzhelian. The first second order sequence presents

three third order sequences and shows a lowstand system tract with a strong erosion between the base of

the sandstone on N, and the coal seam nj below N 4 . a transgressive system tract between n] and O,, a

highstand system tract between O, and the base of the sandstone on OJ . The middle and upper

Source: MNHN, Palis

STRATIGRAPHY OF UPPER CARBONIFEROUS AND LOWER PERMIAN IN DONETS BASIN

(2) -(5).

(3) V v (6)-

F,G. 7.— Lithostratigraphy and sequence stratigraphy of the upper GzheUan and Lower p ermian in the Donets Basin. A:

Artemovsk. G: Gorlovka. I: Ivano-darievka. K: Khourakhovka. L: Lisitschansk. O: Otcheremuno. (1) j

actual, b: supposed; (2) Palynological zone; (3) Halite; (4) Sequence boundary; (5) Flooding surface, (6) Maximal

FlG 7 —Uthosnaiteraphie et stratigraphic sequentielle du Gzhelien supirieur et du Permien inferieur dans le basstn du

DoneteA Artemovsk. G : Gorlovka. I . Ivano-darievka. K : Khourakhovka. L : Lislschansk. O : Otcheremuno (1)

Catcaire, a - observe, b : suppose ; (2, biozone palynologique ; (3) Halite ; (4) limtte de sequence : (5, surface

d'inundation ; (6) surface d'inondation maximale.

Kasimovian can present gaps in the Donets Basin. The second second order sequence presents four third

order sequences and shows a lowstand system tract with a strong erosion between the base ot the

sandstone on 0 4 5 and o^ below 0 6 a transgressive system tract between o 5 2 and O,, a highstand system tract

between O, and the base of the sandstone on Q s .

The Asselian is a second order sequence with a lowstand system tract at the base, a transgressive

system tract between Q„ and S„ a maximal flooding surface in S, known in the Dnieper Basin, a

highstand system tract between S, and TIn the Ural Basin, work in progress shows that the Asselian is

also a second order sequence with a highstand system tract in the third third order sequence. In the

Donets Basin, only the lower Sakmarian is present and continues the Asselian sequence. Artinskian,

Kungurian are absent and only Tatarian is present in the Upper Permian.

DISCUSSION ON THE GEOLOGICAL CONTROLS

The geological controls on the third and fourth order sequences are multiple. They include local and

regional tectonic subsidence and uplift, sedimentary processes, climate and eustasy ot glacial or plate-

tectonics origin. The sequence periodicity does not allow separation ot the glacio or tectonic eustasy tor

ALAIN IZART ET AL.

DNIEPER

DONETS

Source: MNHN. Paris

STRATIGRAPHY OF UPPER CARBONIFEROUS AND LOWER PERMIAN IN DONETS BASIN

the third and fourth order sequences (CARTER et al., 1991) as well as the effects of regional tectonics

(DICKINSON et al., 1994, p. 29). Using the spatial sequence amplitude, it is possible to separate local,

regional and global events. The fourth, third and second order sequences have been identified all over

the Donets Basin. The fourth order sequences are different in the transitional area and the more marine

area. However, it should be possible to follow them and find their equivalents as well as their

continuation in the Donets and Dnieper basins. For the elementary sequences we assume the continuity

of the marine band, limestone becoming a lagoonal claystone westwards, indicating a maximal flooding

surface. However, erosion has destroyed many elementary sequences in the more continental area and it

is illusive to follow them far to the west across the continental area.

IZART & VACHARD (1994) hoped to separate tectonic and eustatic controls in comparing the Moscow

Basin, which shows a little tectonic subsidence and is more sensitive to eustasy, with the Donets Basin

where the tectonic subsidence is more important. Both basins are located on the same plate, the Russian

plate and both basins have an open connection with the Tethys ocean. In the Moscow Basin, BRIAND et

al. (in press) checked the validity of the fourth order sequences described by YABLOKOV et al. (1975).

The same sequences were found in the Moscow and Donets basins. However, this result does not prove

their glacio-eustatic origin because the Moscow basin became subjected to uplift of the Voronezh horst

that separated the Donets Basin and the Moscow Basin. Numerous uplifts of the platform are recorded

for the Moscow Basin during the Bashkirian and the lower Moscovian and regional tectonics could

produce these sequences. However, each third order sequence corresponds to a biozone of fusulinids.

Lithological and biological sequences are identical. Therefore, allocyclic factors including tectonic

subsidence, eustasy and climate produce these evolutions.

TUCKER et al. (1993, p. 399) showed that during the "icehouse" time, when an icecap covered the

pole, the third order sequences present a smaller amplitude than the fourth order sequence. During the

Carboniferous, a polar ice cap existed on Gondwana (VEEVERS & POWELL, 1987; CROWELL, 1995) and

could explain these fourth order sequences. The study of the sequence cyclicity by Walsh analysis

(BEAUCHAMP, 1984 ) and comparison with Milankovitch periodicities (De BOER & SMITH, 1994)

allows to test this hypothesis. In that case, a better accuracy of time by the radiochronology is needed to

calibrate the sequences of the Donets basin.

The drift of the Pangea into equatorial latitude during Carboniferous and tropical latitude near the

Carboniferous / Permian boundary (SCOTESE & McKERROW, 1990; SCOTESE & LANGFORD, 1995)

explains the differences in climate (wet / dry), lithology (coal / evaporites, red beds) and plants

(hygrophilous / xerophylous) between Carboniferous and Permian. The disappearance of coals, the

appearance of red beds and evaporites in the Donets Basin are explained by the increase of dryness and

establishment of monsoons induced by the perfect symmetry of continents forming Pangea

approximately at the equator (BESLY, 1987; KUTZBACH & GALLIMORE, 1989; CECIL, 1990; BARRON &

Fawcett, 1995; Parrish 1995).

FIG 8 — Blocks diagrams of the Dnieper-Donets rift during the genesis of a fourth order sequence in the Kasimovian and the

Gzhe^an time A: Uplift of the Voronezh and Ukrainian horsts and / or global eustatic fall, erosion and fluvial

deposition in the Dnieper-Donets graben. This period corresponds to a tectonic and/or eustatic owstand ‘ ra ct. B.

Tectonic subsidence in the Dnieper-Donets graben and / or global eustatic rise with flooding that stops the fluvial

deposits up-stream and allows the settlement of large swamp (coal). This period corresponds to a tectonic andI /or

eustatic transgressive system tract. C: Tectonic subsidence in the Dnieper-Donets graben and / or global eustatic rise

with a maximal flooding and deposits of the marine limestones and claystones. This period corresponds to a tectonic

and / or eustatic maximal flooding period. D: Tectonic subsidence in the Dnieper-Donets graben and / or filling by the

deltas At first, the environment is marine and then, gradually becomes lagoonal or lacustrine. This period corresponds

F,0 S.'° Bte ior, * ,a g M,e Sun. Su «*. p.nSuu, I.

Kasimovien It le Gzhelien. A : Soulevement des horsts du Voronezh et de l Ukraine et / ou chute eustatique S^baje

erosTon et dlvdt fluvianle dans le graben da Dnieper-Donetz. Celle periode correspond an cortege de has niveau

lectoniaue el / ou eustatique. B : Subsidence tectonique dans le graben du Dnieper-Donetz et / ou montee eustatique

globTavec I’inondation qui bloque le depot fluvialHe en a,non, e, perrnel / ^gu^J,

erande surface Cette periode correspond a an cortege transgresstf tectomque et / ou eustatique. C . Subsidence

lectoniaue dans le graben du Dnieper-Donetz et / ou montee eustatique globale avec une mondation maximale et des

depots de calcaires'et argilites marines. Cette periode correspond a la periode dinondation max,male tectomque et/ou

eustaUque D Subsidence tectomque dans le graben du Dnieper-Donetz e, / <>,< rempltssage par les deltas

L’environnement es, tout d'abord marin. e, puis graduellement ,I devtent laguna,re ou locus,re. Cette penode

correspond du cortege de haul niveau tectonique et / ou eustatique.

Source . MNHN. Pans

ALAIN IZART ET AL.

Tectonic events at the scale of the Russian plate were the cause of these fourth order sequences with

a superposition of glacio-eustatic oscillations. Figure 8 represents block diagrams of the Dnieper-Donets

rift during a fourth order sequence in Late Carboniferous time. Each sequence presents four periods:

— the uplift of the Voronezh and Ukrainian horsts and / or the global eustatic fall, the erosion and the

fluvial deposition in the Dnieper-Donets graben. This period corresponds to a tectonic and / or eustatic

lowstand system tract.

— the tectonic subsidence in the Dnieper-Donets graben and / or global eustatic rise with a flooding

that stops the fluvial deposits up-stream and allows the settlement of large swamp (coal seams). This

period corresponds to a tectonic and / or eustatic transgressive system tract.

— the tectonic subsidence in the Dnieper-Donets graben and / or global eustatic rise with a maximal

flooding and deposits of marine limestones and claystones. This period corresponds to a tectonic and / or

eustatic maximal flooding period.

— the tectonic subsidence in the Dnieper-Donets graben and / or the filling by deltas. At first, the

environment is marine and then, gradually becomes lagoonal or lacustrine. This period corresponds to a

tectonic and / or eustatic highstand system tract.

An uplift or a tectonic inversion period occur between lower Sakmarian and Tatarian and between

Tatarian and Lower Triassic.

CONCLUSION

New results concern the stratigraphy and the sequence stratigraphy of the Kasimovian, Gzhelian and

Lower Permian in the Donets Basin and the geological controls in this basin. The Kasimovian

corresponds to the lower Stephanian of the western Europe, the Gzhelian to the upper Stephanian and

the lowermost Autunian, the Asselian to the uppermost Autunian. The Kasimovian, Gzhelian and

Asselian are second order sequences and the fourth order sequences obscure the third order sequences in

the Donets basin. This fact was yet known in the Moscow basin, the western Europe and the USA

(IZART & VACHARD, 1994). Nevertheless, it is possible to distinguish three third order sequences

equivalent to the lower, middle and upper Kasimovian, four third order sequences equivalent to the

Gzhelian and three third order sequences equivalent to the Asselian, which can be further subdivided

into fourth order sequences. The sequences are controlled by the regional tectonic subsidence in the

graben and the uplift of the horsts and / or by eustasy with either a plate-tectonic or a glacial origin.

Each third order sequence corresponds to a biozone of fusulinids. Lithological and biological sequences

are identical. The thickness and tectonic subsidence decrease gradually between the Moscovian and

Sakmarian. Hiatus exist betwen the Sakmarian and Tatarian and between Tatarian and Lower Triassic,

corresponding to an uplift or a tectonic inversion period. The drift of the Pangea into equatorial latitude

during Carboniferous and tropical latitude near the Carboniferous / Permian boundary explains the

differences in climate, lithology and plants between Carboniferous and Permian.

ACKNOWLEDGEMENTS

We thank Professor STSCHEGOLEV, all the participants of the field mission during year 1994 in the

Donets basin and N. ZHIKALIAK of the Donbass geological survey in Artemovsk. This study was carried

out under a French-Ukrainian project financed by the Peri-Tethys program and BRGM. This note is a

contribution of the IGCP 343, the Peri-Tethys program and of the UMR-CNRS 7566 G2R.

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Source: MNHN. Paris

STRATIGRAPHY OF UPPER CARBONIFEROUS AND LOWER PERMIAN IN DONETS BASIN

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Stratigraphic correlations of the Upper Permian and

Triassic beds from the Volga-Ural and Cis-Caspian

Edward A. MOLOSTOVSKY, Ija I. MOLOSTOVSKAYA

& Maxim G. Min/KH

Saratov State University, Moskovskaya str.. 161,410750 Saratov. Russia

ABSTRACT

Correlation of the Upper Permian and Triassic deposits from the East of the Russian Plate, Cis-Ural Trough and the North

of the Cis-Caspian Depression was performed on the basis of paleontologic. paleomagnetic and lithologic-facies data. Main

sedimentation stages were distinguished, the influence of the geologic and climatic factors on sedimentation processes was

analyzed. Of main importance in this respect were: hydrochemical regime of the Early Permian residual basin, transgressions of

the Boreal Sea during the Kazanian, the Tethys transgressions in the Triassic and increased terrigenous drift from the region of

the folded Urals during its tectonic activizations. Stratigraphic units were ranked according the scales and characters of

accompanying geologic events.

RESUME

Correlations stratigraphiques dans le Permien superieur et le Trias des regions Volga-Oural et Cis-Caspienne.

Les correlations des depots du Permien superieur et du Trias de l’Est de la Plate-forme russe, du Cis-Oural et du Nord de la

Cis-Caspienne ont ete realisees sur la base de donnees paleontologiques, paleomagnetiques et faciologiques. Les principales

etapes de la sedimentation sont distinguees, V influence des facteurs geologiques et climatologiques sur les processus de

sedimentation est analysee. De ce point de vue. les evenements les plus importants sont le regime hydrochimique du bassin

residue! du Permien inferieur, les trangressions de la Mer Boreale durant le Kazanien. les transgressions de la Tethys au Trias et

Paugmentation des apports detritiques depuis la region plissee de 1’Oural lors de son activite tectonique. Les unites

stratigraphiques sont organisees selon les echelles et les caracteres des evenements geologiques.

INTRODUCTION

The South Cis-Ural, Lower Volga and North Cis-Caspian regions include a vast area between the

Volga River and the Ural Range, with the conventional southern border along the latitude of Inder and

Baskunchak lakes (Astrakhan region) and the northern one - in the lower reaches of the Belaya River

(Bashkiria). The boundaries of three major geostructures meet in this region: those of the Russian Plate,

Cis-Ural Trough and Cis-Caspian Depression (Fig. 1). The main features of the Permo-Triassic

sedimentation were determined by the development of these structures.

MOLOSTOVSKY, E.A., Molostovskaya. 1.1. & MlNlKH, M.G., 1998.— Stratigraphic correlations of the Upper Permian and

Triassic beds from the Volga-Ural and Cis-Caspian. In: S. Crasquin-Soleau & E. Barrier (eds), Peri-Tcthys Memoir 3:

stratigraphy and evolution of Peri-Tethyan platforms. Mem. Mus. natn. Hist. nat.. Ill : 35-44. Paris ISBN : 2-85653-512-7.

Source: MNHN. Paris

EDWARD A. MOLOSTOVSKY ETAL.

From a stratigraphic point of view, the Volga-Ural region is referential for the northern margin of the

Tethys. Type sections for the majority of the continental Upper Permian and Triassic units are

concentrated here; stratigraphic sequences of the index faunal groupings are studied thoroughly; a

detailed magnetic zonality scheme has been produced.

The present paper is based upon geological, paleontological and paleomagnetic materials gained by

the authors (MlNlKH. 1972, 1977, 1995; MlNIKH & MlNlKH, 1985, I995 ;"Mc>LOSTOVSKAYA, 1987,

1993; MOLOSTOVSKY, 1983, 1995; BLOM et aL, 1982; LIPATOVA et al., 1985) and on the research

results obtained by a substantial number of geologists (GORSKY & GUSEVA (eds) 1990; SHISHKIN &

OCHEV, 1992; TVERDOKHLEBOVA, 1991; ZAMARENOV et al, 1969).

Facies variability of the red-bed sequences makes it difficult to correlate them according to lithologic

features, which are of local importance. Regional and remote correlations are usually carried out

according to the index fossil complexes; the most accurate results in some stratigraphic intervals are

available by means of combined paleomagnetic and paleontologic definitions.

Ostracods are the leading stratigraphic group for the Upper Permian red-bed formation; they allow to

divide and correlate the sequences at the stage and horizon levels (GORSKY & GUSEVA eds, 1990;

MOLOSTOVSKAYA, 1987, 1993). Terrestrial vertebrates make it possible to recognize only the upper

Tatarian substage, characterized by the Pareiasaur-Theriodont fauna. The more ancient. Dinocephalian

fauna is common for the Ufimian and Kazanian stages, and for the lower Tatarian substage as well

(Gorsky & Guseva eds, 1990).

Tetrapods and lungfish are of guiding stratigraphic importance for the Lower and Middle Triassic,

palynoflora is important for the upper section (MlNIKH & MlNlKH, 1985, 1995; Lipatova, MlNIKH,

MOLOSTOVSKY et al., 1985; SHISHKIN & OCHEV, 1992). The paleomagnetic data become

stratigraphic-ally important in the upper Tatarian substage of the Permian and in the Triassic, where

repeated changes of magnetic zonality are recorded. The Ufimian, Kazanian and lower Tatarian beds

were formed during the stable period of reverse field (R-Kiama) and are not divisible according to

paleomagnetic characteristics (MOLOSTOVSKY, 1983).

The main regional and local units from various structural-facies zones are presented in figures 2 and

3 together with the guide fossil complexes and schemes of magnetic zonality. The figures were based

upon the unified stratigraphic schemes for the Russian Plate, Cis-Ural Trough and Cis-Caspian

Depression (GORSKY & GUSEVA eds, 1990; ZHAMOYDA. Lipatova & Romanovskaya (eds), 1982).

STRATIGRAPHIC CORRELATIONS

Upper Permian

In the most part ot the territory considered, all the stages of the Upper Permian are represented in the

sections.

The Ufimian stage in the south-east of the Russian Plate is composed of lacustrine-lagoonal deposits:

red. more rarely grey clays, siltstones, sandstones, limestones, dolomites with anhydrite and gypsum

interlayers (50-200 m thick). Mostly redstones, gypsinate in the lower portion of the section, occur to the

east, where the Platform passes to the Trough (up to 190 m thick). Sandstones and coarse-pebble

conglomerates prevail in the sections from the eastern part of the Trough (up to 700 m thick). In the

north-eastern part of the Cis-Caspian Depression, the Akshatskaya formation (750 m thick), composed

of alternating red and grey argillites and sandstones with anhydrite interlayers, is referred to the Ufimian

stage.

The Ufimian sections of various facies are correlated according to ostracods belonging to the zonal

complex of Paleodarwinula abunda (Mandelstam) and miospores of the zonal complex H2-

Granizospora vulgaris (Naumova).

The Kazanian stage is subdivided into the lower and upper substages.

The lower Kazanian substage within the south-east of the Russian Plate, the western zone of the Cis-

Ural Trough and the north-western part of the Cis-Caspian Depression, comprises various grey marine

rocks up to 120-150 m thick. In the sections from the south-east of the Russian Plate, the lower

Kazanian sequence is carbonate-halogenic; limestones and dolomites dominate, sandstones, halite and

Source: MNHN. Paris

UPPER PERMIAN AND TRIASSIC OF VOLGA-URAL AND CIS-CASPIAN

anhydrites are less common. According to the ratio of various lithotypes, the substage there comprises

the Bajtuganskaya, Kamyshlinskaya and Barbashinskaya members (GORSKY & GUSEVA eds, 1990). In

the western part of the Trough, clay-aleurolitic deposits with occasional interlayers of bioaccumulated

limestones are common (70-100 m thick). In the axial zones, marine facies give way to lacustrine-

lagoonal grey sandstones, siltstones and argilites containing charred plant detritus and rare marl

interlayers (up to 600 m thick). In the eastern part of the Cis-Ural Trough, accumulations from saliferous

lagoons occur: dolomites, anhydrites, gypsums with halitic interlayers (KULEVA, 1975).

The marine beds contain numerous remains of toraminiters, brachiopods, fish, leaty flora and

miospores of the zonal complex J1 - Striatopodocarpites tojmensis Sedova. The palynocomplexes from

EDWARD A. MOLOSTOVSKY ETAL

THE SOUTH-EAST OF

THE RUSSIAN PLATE

(LEFT BANK VOLGA REGION

IN KUJBYSHEV AREA,

ORENBURG CIS-URALS)

CONTACT ZONE OF THE

RUSSIAN PLATFORM ANO

THE CIS URAL

MARGINAL TROUGH

(ORENBURG CIS URALS)

AXIAL AND EASTERN

ZONES OF THE CIS URAL

MARGINAL TROUGH

(ORENBURG CIS-URALS)

NORTH EASTERN BORDE-

RING ZONE OF THE CIS-

CASPIAN DEPRESSION

(AKTYBINSK CIS URALS)

<$H- a <3>i b

Fig. 2.— Correlation of the Upper Permian regional stratigraphic schemes.

Symbol for figures 2 and 3: a. tetrapods; b. fishes; c. bivalves: d, ammonites; e, brachiopods; f, ostracods; g,

conchostracans: h. foraminifers; i. harofites; j, miospores; 1, flora.

Guide fauna and flora groups in the Permian stratigraphy: Tetrapods: 1, Dinocephalia, 2, Chroniosaurus dongusensis

Tverdochlebova; 3, Chroniosuchus uralensis Tverdochlebova. Bivalvia: 4, Palaeomutela vjatkensis Gusev. Ostracod: 1,

Paleodarwinula abunda (Mandelstam); 2, Amphissites tscherdynzevi Posner; 3 ,Paleodarwinula fainae (Belousova); 4,

Paleodarwinula fragilifonnis (Kashevarova), Suchonellina futshiki (Kashevarova), Suchonellina trapesoidci (Sharapova).

Miospores: H. Granizospora vulgaris Naumova; Jl. Striatopodocarpites tojmensis Sedova; J2, Lueckisporites virkkia

Potonie & Kremp.

Fig. 2.— Correlation des schemas stratigraphiques regionaux du Permien superieur.

Symboles pour les figures 2 et 3: a. tetrapodes; b, poissons; c. bivalves; d, ammonites; e, brachiopodes; f ostracodes;

g. conchostraces; h. foraminiferes; i, harofites; j, miospores; l.flore.

Grounes guides de la faune et la flore du Permien superieur. Tetrapodes : 1. Dinocephalia ; 2, Chroniosaurus

dongusensis Tverdochlebova ; 3. Chroniosuchus uralensis Tverdochlebova. Bivalves : 4, Palaeomutela vjatkensis Gusev.

Ostracodes : I. Paleodarwinula abunda < Mandelstam ); 2, Amphissites tscherdynzevi Posner ; 3, Paleodarwinula fainae

(Belousova ) ; 4. Paleodarwinula fragiliformis ( Kashevarova ), Suchonellina futshiki (Kashevarova). Suchonellina

trapesoida (Sharapova). Miospores : H , Granizospora vulgaris Naumova \J1 , Striatopodocarpites tojmensisSeYfova ; J2,

Lueckisporites virkkia Potonie & Kremp.

the Jl zone serve as the most reliable correlatives in comparing the lower Kazanian non-marine beds

with each other and with the marine analogues.

The upper Kazanian substage within the Platform is composed mostly of saliferous-lagoon

sediments, subdivided from bottom to top into the hydrochemical (10-200 m thick), Sosnovskaya (30-

125 m thick) and Sokskaya (up to 100 m thick) members. The lowermost member is composed of

anhydrites and, to a lesser extent, of rock salt. The Sosnovskaya member is represented by anhydrites,

dolomites and limestones. Red sandstones with gypsum lenses and nests prevail in the Sokskaya

member. In the zone of the Platform/Trough contact and within the trough, the upper Kazanian sequence

consists mainly of red siltstones and clays, interlayered with sandstones, marls and limestones (up to 700

m thick).

Source: MNHN. Paris

UPPER PERMIAN AND TRLASSIC OF VOLGA-URAL AND CIS-CASPIAN

Ostracod zonal guide fossils Darwinula fainae (Belousova) and miospores of the palynozone J2 -

Lueckisporites virkkia Potonie & Kremp are common for all the rocks of diverse geneses. As a whole,

the non-marine tacies present in the Kazanian stage, are characterized by guide fossil bivalves

Palaeomutela umbonata (Fischer).

Recognition and division of the Kazanian deposits within the Cis-Caspian Depression is difficult due

to poor paleontological description and insufficient data, especially from its western part. In the eastern

side zone, the Kazanian stage contains the Blagodarnenskaya formation (750 m thick), built of

alternating argilites, siltstones and sandstones interlayered by limestones. Red beds prevail in the upper

part of the section; the lower half is dominated by grey rocks containing fossil conifer pollen,

Striatopodocarpites , characteristic of the lower Kazanian substage from the Cis-Urals.

The Tatarian stage is divided into the lower and upper substages.

The lower Tatarian substage within the Urzhum Platform sections is subdivided from below upwards

into the Bolshekinelskaya (50-100 m thick) and Amanakskaya (45-90 m thick) formations. The

Bolshekinelskaya formation is composed of cross-bedded alluvial sandstones interlayered by red clays;

the Amanakskaya one of clays and siltstones with rare marl and limestone layers.

In the zone of the Platform/Trough contact, the Salmysh sandstones and conglomerates (90-100 m

thick) and the Grebenskaya clay-siltstones (80-100 m thick) are analogous to those mentioned above. In

the Cis-Ural Trough, the Urhzum horizon is includes a thick (up to 800 m) red-bed sequence of

sandstones, siltstones, conglomerates and clays interlayered with nodular limestones. In the north¬

eastern part of the Cis-Caspian Depression, the lower substage includes the Tuketskaya and

Aktyubinskaya formations. The Tuketskaya one (up to 360 m thick) is composed of red sandstones,

conglomerates, siltstones and clays; the Aktyubinskaya one (660 m thick) is of finer terrigenous

composition. The substage is characterized by zonal guide fossils of ostracods Paleodarwinula

fragiliformis (Kashevarova) and bivalves Palaeomutela vjatkensis Gusev.

The upper Tatarian substage is subdivided into the Severodvinsky and Vyatsky horizons,

everywhere, except for the north-east of the Cis-Caspian Depression. In this part, the red-bed

Rodnikovskaya formation composed of siltstones and sandstones, interlayered with limestones (820 m

thick), corresponds to the undivided upper substage. In the Platform regions, the Severodvinsky horizon

corresponds to the red-bed Malokinelskaya formation, composed of alternating clays, siltstones and

sandstones interlayered with marls and limestones (up to 170 m thick). In the zone of the

Platform/Trough contact, only terrigenous deposits occur: sandstones, conglomerates, siltstones (up to

100 m thick). In the Cis-Ural Trough, the Severodvinsky horizon (up to 700 m thick) has variable

carbonate-terrigenous composition. The Severodvinsky horizon is distinguished on the basis of the zonal

guide fossils of ostracods Suchonellina futschiki (Kashevarova) and the tetrapods Chroniosaurus

dongusensis Tverdochlebova (TVERDOKHLEBOVA, 1991). In the paleomagnetic scale, zones N1P2 and

R2P2 are equivalent to the Severodvinsky horizon (MOLOSTOVSKY, 1983). The Vyatsky horizon in the

Platform, is represented by the Kutulukskaya formation composed of red and variegated clays,

siltstones, marls, limestones with gravelstone and sandstone lenses (up to 100 m thick). In the Cis-Ural

Trough, the horizon section is dominated by brown clays, siltstones and sandstones; grey and brown

marls and limestones are less common (up to 460 m thick).

The Vyatkian deposits contain numerous remains of bivalves, conchostracans, charophytes and

fishes. Ostracods Suchonellina fragiloides (Zekina), Suchonella typica Spisharsky and tetrapods

Chroniosuchus uralensis Tverdochlebova are guide of stratigraphic importance. In the paleomagnetic

scale, zones N2P2 and R3P2 correlate with Vyatkian horizon.

LOWER TR1ASSIC

In the South Cis-Urals, the Lower Triassic is represented by the Blumental group and Petropavlovka

formation.

The Blumental group includes the beds of three major cycles of alluvial sedimentation,

corresponding (from bottom to top) to Kopanskaya (up to 590 m), Staritskaya (up to 370 m) and

Kzylsajskaya (up to 330 m) formations. Each one of those is composed of three to seven rhythmic

members, beginning with cross-bedded sandstones with conglomerate lenses. The upper parts of the

EDWARD A. MOLOSTOVSKY ETAL.

Fig. 3.— Correlation of the Triassic regional stratigraphic schemes.

Zonal guide fossil in the Triassic stratigraphy:

Tetrapods: 1, Tupilakosaurus ; 2. Benthosuchus ; 3, Wetlugasaurus ; 4, Parotosuchus ; 5, Eryosuchus\ 6, Mastodonsaurus.

Fishes: I, Gnathorhiza triassica Minikh: 2, Ceratodus multicristatus Vorobjeva; 3. Ceratodus gracilis Vorobjeva; 4,

Ceratodus bucobaensis Minikh. Ostracods: 1, Darwinula ovalis Glebova; 2, Darwinula longissima Belousova; 3,

Darwinula lauia Schneider; 4, Lutkevichinella brutlanae Schneider: 5, Glorianella inderica\ 6, Pulviella aralsorica

Schneider: 7, Gemanella schweyeri ; 8, the upper Triassic biozone. Conchostracans: 1. Vertexia tauricornis Lutkevich;

2, lower Baskunchak biozone; 3, Tananyk bio-assemblage; 4, Bogdo bio-assemblage; 5, upper Baskunchak bio¬

assemblage; 6, the Middle Triassic; 7, the upper Triassic biozone. Miospores: 1-8. Miospore biozone (MBZ).

FlG. 3 .— Correlations des schemas stratigraphiques regionaux du Trias.

Fossiles guides de la stratigraphie du Trias.

Tetrapodes : I, Tupilakosaurus ; 2. Benthosuchus ; 3, Wetlugasaurus ; 4 . Parotosuchus ; 5, Eryosuchus ;

6, Mastodonsaurus. Poissons : I, Gnathorhiza triassica Minikh ; 2, Ceratodus multicristatus Vorobjeva \3, Ceratodus

gracilis Vorobjeva ; 4, Ceratodus bucobaensis Minikh. Ostracodes : 1, Darwinula ovalis Glebova ; 2, Darwinula

longissima Belousova ; 3, Darwinula lauta Schneider ; 4, Lutkevichinella bruttanae Schneider ; 5, Glorianella inderica ;

6, Pulviella aralsorica Schneider ; 7, Gemanella schweyeri; 8. Biozone du Trias superieur. Conchostraces : l , Vertexia

tauricornis Lutkevich ; 2. Biozone du Baskunchak inferieur ; 3, Bio-assemblage de Tananyk ; 4, Bio-assemblage de

Bogdo ; 5. Bio-assemblage du Baskunchak superieur; 6, Le Trias moyen ; 7. Biozone du Trias superieur.

Miospores : 1-8, biozones a miospores (MBZ).

members are generally clayey. Less thick analogues of the suites mentioned, are recognized in the

sections from the left bank of the Volga.

Some remains of neorachitomous labyrinthodonts, Tupilakosaurus sp., occur in the Kopanskaya

formation; Bentosuchus cf. sushkini (Efremov) in the Staritskaya formation; Wetlugasaurus angustifrons

Riabinin in the Kzylsajskaya formation. Remains of the dipnoan Gnathorhiza triassica Minikh are

known from two upper formations. In the paleomagnetic scale, zones N1T1, R1TI and partially R2T1

correspond to the Blumental group. According to the neorachitomous tetrapod fauna and magnetic

zonality, the Blumental group correlates with the Vetluga supergroup from the Russian Plate and the

Indian stage and the lower Olenekian substage from the Boreal region (SHISHKIN & OCHEV, 1992 ).

Source: MNHN. Paris

UPPER PERMIAN AND TRIASSIC OF VOLGA-URAL AND CIS-CASPIAN

The Petropavlovskaya formation is composed of a number of rhythmic members opening with red-

brown sandstones of alluvial-deltaic type and completed with lacustrine-floodplain siltstones and clays.

The formation contains the parotosuchian tetrapod fauna and the Yarenian ichthyofaunal complex with

Ceratodus multicristatus Vorobjeva. The lower part of the formation is normally magnetized and

belongs to the zone N2T1, the upper part of the suite corresponds to the magnetozone R1T1.

In the east ot the Cis-Ural Trough, the Lower Triassic includes the Girjalskaya member of sandstones

and conglomerates of alluvial-proluvial origin (up to 420 m). In the north-eastern part of the Cis-Caspian

Depression the red-bed terrigenous Pericaspian series ot alluvial-deltaic type corresponds to the Lower

Triassic (up to 1600 m); it consists ol the Ershovskaya and Zhulidovskaya formations with peculiar

freshwater zonal guide fossils of ostracods, conchostracans and charophytes.

In the south-western part of the depression, in the lower section, the following units are recognized

(upwards): chiefly red-bed alluvial-deltaic, silto-arenaceous Bugrinskaya formation (up to 600 m), sand

sequence (up to 55 m), sand-conglomerate sequence (up to 45 m) and silto-clayey (variegated in its

upper part) Akhtubinskaya formation (up to 110 m), formed in shallow-water, littoral settings. These are

overlain by the marine, grey-coloured clay-carbonate Bogdinskaya formation (up to 330 m), in its turn

overlain by the variegated, sandy-clayey Yenotayevskaya formation (up to 200 m), accumulated under

the conditions ol near-shore plains. According to the faunas of ostracods and conchostracans, the

Bugrinskaya member is referred to the Ershovs formation; the Akhtubinskaya, Bogdinskaya and

Yenotayevskaya members form the Baskunchakskaya formation with its stratotype in Bolshoye Bogdo

Mt. The Bogdinskaya member contains remains ol the Parotosuchus fauna, the Yarenian ichthyofauna

with Ceratodus multicristatus Vorobjeva, and ammonites ( Tirolites cassianus Quenstedt) from the

Olenekian Columbites zone.

In the paleomagnetic scheme the sand sequence and sand-conglomerate sequence corresponds to the

zone R1T1, the Akhtubinskaya and preserved part of the Bogdinskaya members to the zone N2T1.

Middle Triassic

In the southern part of the Cis-Ural Trough, the Middle Triassic corresponds to the Donguzskaya,

Bukobajskaya and the lowermost part of the Surokajskaya formation.

The Donguzskaya formation consists of variegated siltstones and clays of lacustrine origin with

interlayers and lenses of grey cross-bedded sandstones (up to 360 m). The formation is characterized by

tetrapods of Eryosuchus fauna, the Donguz Middle Triassic ichthyofauna with Ceratodus gracilis

Vorobjeva, and the middle Anisian and early Ladinian guide fossil miospores (Fig. 3).

The Bukobajskaya formation consists of variegated siltstones and clays rich in plant remains; they

alternate rhythmically with lenses of grey, cross-bedded sandstones (up to 600 m). The formation is

characterized by tetrapods of Mastodotisaurus fauna, the Bukobajskaya zone of the Middle Triassic

ichthyofauna with Ceratodus bukobaensis Minikh. According to its guide fossil flora and miospores, the

Bukobajskaya formation correlates with the late Ladinian.

In the Cis-Caspian Depression, the Middle Triassic is recognized as the Akmajskaya series, which

includes (from bottom to top) the Eltonskaya, Inderskaya and Masteksajskaya formations.

The Eltonskaya formation consists of littoral-marine variegated grey clays and limestones (up to 300

m) with guide fossil charophytes, ostracods and miospores, which allow to refer the formation to the

Anisian. In the vicinity of Inder Lake, the formation contains occasional dipnoans Ceratodus gracilis

Vorobjeva, from the Donguz zone (Fig. 3).

The Inderskaya formation (up to 220 m) is represented by marine grey clays and limestones with the

Ladinian guide fossil charophytes, ostracods and miospore complexes. In the uppermost part of the

formation, there are some remains of the Bukobajskaya dipnoan and the mastodonsaur tetrapod faunas.

The Masteksajskaya formation (up to 210 m) consists of dark-grey clays interlayered by siltstones

and sandstones. Marine ostracods from the Gemanella zone and the late Ladinian miospores are

characteristic of this formation.

Upper Triassic

The Surakajskaya formation (up to 300 m) corresponds to the Upper Triassic in the South Cis-Urals;

it consists of lacustrine-marsh grey-coloured clays, siltstones and sandstones. According to the

EDWARD A. MOLOSTOVSKY ET AL.

paly nologic data, the formation is composed of three biozones, MBZ-6,7,8, corresponding to the

Camian, Norian and Rhaetian stages, respectively (Fig. 3).

In the Peri-Caspian Depression, the Upper Triassic corresponds to the continental grey-coloured-

variegated coal-bearing Aralsorskaya group (Fig. 3), comprising (from bottom to top) the

Akmamykskaya (up to 425 m), Khobdinskaya (up to 380 m) and Kusankudukskaya (up to 300 m)

formations. Each of them represents a sedimentary cycle starting with sandstones and ending with clays.

All the formations include the Late Triassic zonal guide fossils of conchostracans and ostracods, and are

correlated according to miospores with the Carnian, Norian and Rhaetian stages of the general scale,

respectively.

PALEOGEOGRAPHIC RECONSTRUCTIONS

In the Upper Permian sections from the east of the Russian Plate and the west of the Trough, three

lithologic-stratigraphic complexes are recognized: the Ufimian continental red-bed, the lower Kazanian

marine grey-coloured and the upper Kazanian-Tatarian red-bed ones. A similarly composed three-

member sequence is recognized in the east of the Cis-Caspian Depression and the central zone of the

Cis-Ural Trough. The lower Kazanian marine formations, however, are replaced there by the coalified

grey-coloured sequence of lacustrine-lagoonal and possibly littoral-marine sediments.

A certain continuity may be recognized in the spatial arrangement of the facies in the Ufimian and

Kazanian sedimentation fields from the Russian Plate and the Cis-Ural Trough. This continuity consists

in preferential concentration of the salt-bearing and carbonate sediments within the western regions of

the Cis-Urals, and their gradual substitution by the terrigenous facies towards the folded Urals. It is

worthy of note that the tendency is retained in the lower Kazanian marine basin that has drastically

changed the sedimentation settings in the region.

For the Ufimian and upper Kazanian red-bed deposits, such continuity is caused by similar

paleogeographic conditions: they were accumulating in shallow relict reservoirs, inherited from the

Kungurian and lower Kazanian basins, relatively remote from the Ural sourcelands. The zones

influenced by the latter ones, were characterized by dominating regime of terrigenous sedimentation. In

the early Kazanian, similar arrangement of the facies was probably caused by uplifting mineralized

waters from the westerly marine basin. On the whole, in all the territories considered, arid continental

conditions were characteristic of the Late Permian sedimentation cycle. In the early Kazanian they were

interrupted by a short-term boreal transgression accompanied by general humidization of the climate.

During the whole of the Kazanian age, the character of sedimentation all over the Volga-Ural and Cis-

Caspian regions was to a variable extent influenced by the Boreal Sea transgression.

With the start of the Triassic sedimentation cycle, independent formational sequences came into

being in the Cis-Ural Trough and Cis-Caspian Depression. In the Cis-Caspian, the three-member series

consisted of red-bed terrigenous, grey-coloured marine and coal-bearing lacustine-marsh formations. A

three-member formational series appeared in the South Cis-Urals, as well: alluvial-proluvial, lacustrine

and lacustrine-marsh coalified structures.

In the Early Triassic, the paleogeographic settings were arid everywhere, but the sedimentation

processes were quite autonomous in every region. The relatively limited Olenekian transgression has not

exerted any appreciable influence upon the continental sedimentation in the Cis-Urals, but the dynamic

uplift of the Bogdinskian Sea has probably promoted the appearance of the deltaic facies in the

Petropavlovskaya formation (SHISHKIN et al. 9 1995).

The inter-regional geologic links became more evident in the Middle Triassic. The development of

extensive Tethys transgression has coincided in time with general tectonic stabilization of the Urals,

disappearance of the Lower Triassic alluvial-proluvial plain and setting of lacustrine sedimentation

under more and more humid climatic conditions.

In the Late Triassic, similar landscape-climatic conditions have set in the vast areas of the South Cis-

Urals, Trans-Urals and Cis-Caspian. which resulted in two associations of formations coming together;

under the conditions of humid landscapes, a coalified terrigenous formation started to accumulate. It is

represented by the coal-bearing Chelyabinskaya group in the middle Trans-Urals, Surakajskaya

formation in the South Cis-Urals and the Aralsorskaya group in the Eastern Cis-Caspian.

Source: MNHN. Paris

UPPER PERMIAN AND TRIASSIC OF VOLGA-URAL AND CIS-CASPIAN

CONCLUSION

The Late Permian and Triassic sedimentation was, on the whole, of arid character; it was as late as in

the Late Triassic that lacustrine-marsh coal-bearing sediments started being generated in the vast plains

of the Cis-Caspian and South Cis-Urals as the result of global increase of humidity.

Among the most important regional geologic factors influencing the formation of sedimentary rock

complexes in the Upper Permian, were: residual salt-generating basins of the Lower Permian and the

vast Kazanian transgression of the Boreal Sea, involving the whole of the eastern Russian Plate, the west

of the Cis-Ural Trough and the north of the Cis-Caspian Depression.

In the end of the Early and in the Middle Triassic, sedimentation in the Cis-Caspian area was in many

respects controlled by the Tethys eustatic oscillations. The peculiar character of the Late Permian and

Early Triassic sedimentation within the Cis-Ural marginal trough was determined by intensive

terrigenous drift from the regions of the folded Urals.

The materials available show that the stratigraphic boundaries of practically all the regional units of

the Upper Permian and Triassic from the Volga-Ural and Cis-Caspian regions, are of event character,

but record geologic events of diverse types and levels. This fact makes it possible to carry out research

along various trends of geologic correlations: from detailed stratigraphic comparisons to comparative

analyses of the large-scale reconstructions in paleogeography and geodynamics.

Of main importance in such research is the choice of reference correlation levels, marked with

coincidence of geodynamic, paleogeographic, biotic, formational and geomagnetic boundaries. As far as

the northern margin of the Tethys is concerned, in the Permo-Triassic part of the scale, the following

boundaries meet this condition: those between the lower and the upper Tatarian substages, between the

Permian and Triassic, and the Lower and Middle Triassic.

The rank and geologic importance of the last two boundaries are well-known. The boundary between

the lower and upper Tatarian substages has not been considered properly, yet, and its current

stratigraphic rank does not correspond to the significant reconstructions that have occurred at this level

in the structure of the continental biota and in geomagnetic polarity regime. This boundary is marked by

large-scale changes in tetrapod fauna: replacement of the dinocephalian grouping by the pareiasaur-

theriodontid one (TVERDOKHLEBOVA, 1991; SHISHKIN & OCHEV, 1992). In the course of the Late

Permian, fast change at the superfamily level occurred in ostracodes (MOLOSTOVSKAYA, 1987, 1993).

There were significant changes in ichthyofauna systematic composition (MlNIKH & MINIKH, 1995). In

paleomagnetism evolution, the regime of stable reverse polarity field was replaced by an epoch of

frequent reversals (MOLOSTOVSKY, 1983).

The above data offer profound grounds for regarding the upper Tatarian as an independent stage,

more important in its extent and paleontologic distinctness than any of the currently known stages of the

Upper Permian.

REFERENCES

Blom, G.I., Lozovsky, V.R., Minikh, M.G. et al., 1982.— The general stratigraphic scale: Moscow, Mezen syneclise, Volga-

Ural anteclise (scheme 1). In: A.I. Zhamoyda, V.V. Lipatova & G.M. Romanovskaya (eds). Decision of the

Interdepartment Stratigraphic Meeting on the Triassic from the East-European Platform (Saratov, 1979). Nedra,

Leningrad: 20-35 (in Russian).

Gorsky, V.P. & Guseva, E.A. (eds), 1990.— Decision of the interdepartment regional stratigraphic meeting on the Middle and

Upper Paleozoic from the Russian Platform: with regional stratigraphic charts. Leningrad, 1988. In: The Permian

system. VSEGEI, Leningrad: 1-47 (in Russian).

KULEVA, G.V., 1975.— The Upper Kazanian and Tatarian Continental Deposits from the Southern Part of Cis-Ural Marginal

Deflection (within the Orenburg and south Bashkirian Cis-Urals). Saratov University Publication, Saratov: 1-164 (in

Russian).

Lipatova, V.V., Minikh, M.G., Molostovsky, E.A. et al ., 1985.— Project of the unified scheme for the Triassic from the

East-European Platform. In: V.V. Lipatova & V.G. OCHEV (eds), Triassic Beds from the East-European Platform.

Saratov University Publication, Saratov: 10-28 (in Russian).

Minikh, M.G., 1972.— Fishes. In: V.V. MENNER & V.V. Lipatova £ds), Stratotype Section of the Lower Triassic Baskunchak

Series from Bolshoye Bogdo Mountain. Saratov University Publication, Saratov: 48-50 (in Russian).

EDWARD A. MOLOSTOVSKY ETAL.

Minikh. M.G., 1977.— Triassic Lungfish from the East of European USSR. Saratov University Publication, Saratov: 1-96 (in

Russian).

Minikh. M.G.. 1995.— Fishes. In: M.A. Shishkin (ed.). Biostratigraphy of the Continental Triassic from the South Cis- Urals.

Nauka, Moscow: 38-56 (in Russian).

Minikh, M.G. & Minikh, A.V.. 1985.— Division of the Triassic deposits from the East-European Platform according to

ichthyofauna. In: V.V. Lipatova & V.G. Ochev (eds), Triassic Beds from the East-European Platform. Saratov

University Publication, Saratov: 44-52 (in Russian).

Minikh, M.G. & Minikh, A.V., 1995.— Division of the Permian and Triassic beds from European Russia according to

ichthyofauna. In: Proceedings of the All-Russia Meeting "Paleontology and Stratigraphy of the Continental Permian

and triassic from Northern Eurasia". PIN-RAS, Moscow: 23 (in Russian).

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ostracods. In: V.l. Artemev & G.I. Barulin (eds). Problems of Stratigraphy of the Paleozoic, Mesozoic and Cenozoic.

Saratov University Publication, Saratov: 47-57 (in Russian).

Molostovskaya, 1.1.. 1997.— Non-marine ostracods and paleobiogeographical distribution in the Late Permian basins in the

east of Russian Plate. In: Contribution to Eurasia Geology. Occasional Publications ESRJ. New Series, South Caroline

University Publication, 9: 95-100.

MOLOSTOVSKY, E.A., 1983.— Paleomagnetic Stratigraphy of the Upper Permian and Triassic from East of European Part of

the USSR. Saratov University Publication. Saratov: 1-168 (in Russian).

MOLOSTOVSKY, E.A., 1995.— Magnetoslratigraphic section. In: M.A. Shishkin (ed.). Biostratigraphy of the Continental

Triassic from the South Cis-Urals. Nauka, Moscow: 129-138 (in Russian).

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tetrapod fauna. Bulletin of New Mexico Museum Natural History, Sciences. 3: 435-437.

Shishkin, M.A., Ochev. V.G.. Tverdokhlebov, V.P. et al. , 1995.— Biostratigraphy of the Triassic of Southern Cis-Urals.

Moscow, Nauka: 1-205 (in Russian).

Tverdokhlebova. G.I., 1991.— Tetrapod remain distribution in the Upper Tatarian substage from the East of the European

Platform. Bulletin of Moscow Naturalist Society , 2: 66-71 (in Russian).

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and formation settings of the Upper Permian and Triassic deposits the eastern-border margin of the Peri Caspian

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Nedra. Moscow: 191-201 (in Russian).

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Source: MNHN, Paris

The Domanikoid fades of the Russian Platform

and basin paleogeography

Valentina S. VISHNEVSKAYA

Institute of the Lithosphere. Russian Academy of Sciences, Staromonetny per.. 22. Moscow 109180, Russia

ABSTRACT

The main purpose of the paper is a comparative analysis of the bituminous carbonate-siliceous deposits of the Devonian

(Domanik) and Jurassic (Domanikoid) rift-type basins from the Russian platform to make their correlation, using an

homogeneous chronometer (radiolarians). It compares the chemical composition and lithological types of sediments in order to

understand the dynamics of subsidence and sedimentation, crustal tectonics of the region and to analyze the relationships

between deep and shallow sedimentary basin evolution, as well as establishing the timing of events and. consequently, to

elucidate more general geological processes such as paleoenvironments. We attempt to trace and to interprete the factors which

controlled the sedimentation of domanikoid siliceous rocks enriched in P and C organic in rift basins, contributing to elaborate

Jurassic paleogeographic maps.

RESUME

Le facies Domanikoid sur la Plate-forme russe et paleogeographie du bassin.

Le but de cet article est de presenter une comparaison entre les depots de carbonates siliceux bitumineux du Devonien

(Domanik) et du Jurassique (Domanikoid) dans les bassins de type rift de la Plate-forme russe pour etablir des correlations en

utilisant un chronometre homogene (radiolaires). On compare la composition chimique et le type sedimentologique des

sediments pour comprendre la dynamique de la subsidence et de la sedimentation, la tectonique crustale de la region et pour

analyser les relations entre revolution des zones peu profondes et les zones de bassin aussi bien que pour etablir la chronologie

des evenements et par consequent pour elucider les processus g£ologiques plus generaux comme les paleoenvironnements. On

essaie de tracer et d’interpreter les facteurs qui conlroient la sedimentation des roches siliceuses du Domanikoid enrichies en P

et C organique dans les bassins de rift, contribuant ainsi a elaborer les cartes paleogeographiques du Jurassique.

INTRODUCTION

The presence of domanikoid bituminous facies enriched in organic matter is very typical of Upper

Jurassic strata of the North Peri-Tethys regions (Pechora, Sysola, Volga-Urals hydrocarbon provinces of

the Russian platform). Similarly to the hydrocarbon Domanik Formation of the same area, the origin of

these source rocks probably resulted from the post-rift subsidence events which took place at the rim of

the Russian platform.

Vishnevskaya, V.S., 1998.— The Domanikoid facies of the Russian Platform and basin paleogeography. In: S. Crasquin-

Soleau & E. Barrier (eds), Peri-Tethys Memoir 3: stratigraphy and evolution of Peri-Tethyan platforms. Mem. Mus. natn.

Hist, nat., 177 : 45-69. Paris ISBN : 2-85653-512-7.

Source: MNHN, Pans

VALENTINA S. VISHNEVSKAYA

The continuations of the largest Riphean aulacogens of the Russian platform are traced under Timan-

Pechora and Volga-Urals basins (Fig. 1). The repeated tectonic splitting along the Timan-Pechora,

Vyatka, Volga and Mid-Russian or Moscow (Fig. 2) depression took place during Middle-Late

Devonian and Middle-Late Jurassic time periods. These rifting processes were probably connected to

global world lithosphere plate reorganization at these times.

FIG. 1.— Location of studied samples from Domanik Horizon

(1). The main Urals fault (2) and faults in basement of

Siberian platform; pre-Paleozoic faults (3).

FlG. I.— Localisation des echantillons etudies dans Vhorizon

Domanik (I). Principale faille de iOural (2) et failles

dans le soubassement de la Plate-forme siberienne ;

failles pre-Paleozoiques (3)

Fig. 2.— Location of investigated sites. I-Il, Devonian

paleorifts: I, Pechora, II, Vyatka; III, Riphean

Mid-Russian aulacogen. Names of sites are given

in legend of figure 3.

FlG. 2 .— Localisation des sites etudies. /-II, Paleorifts

devoniens : I, Pechora, II, Vyatka ; III, aula-

cogene ripheen de Russie centrale. Les noms des

sites sont donnes sur la figure 3.

RIFTING AND GEOLOGICAL SETTING

As known, rift-basins are characterized by low crust thickness caused by mantle elevation and its

evolution, higher heat flow values, high seismicity and steep gravity gradient. Measured from regional

deep seismic sounding throughout the Timan-Pechora basin, the Moho depth reaches 36-40 km in its

west part, 38-42 in its center and 34-36 in its north-eastern part (KOSTIUCHENKO, 1993). From

Kostiuchenko's point of view, the spreading (rift) zone is located in the northern part of Timan-Pechora

basin. The linear magnetic anomalies, the uplifting of the consolidated crust surface with a median

valley, and the magnetic body penetrating the crust interpreted as the magma chamber, make the

spreading zone similar to the mid-oceanic ridge. During the Middle and Late Devonian, complex

systems of rift set up along the eastern and south-eastern margins of the European paleocontinent. They

include the Pripyat-Dnieper-Donets, Peri-Caspian, Volga-Ural, Timan-Pechora, Eastern Barents Sea and

Northern Kara Sea rift zones (NIKISHIN el al. y 1993).

Source . MNHN, Paris

DOMANIKOID FACIES OF THE RUSSIAN PLATFORM 47

Fig. 3.— Distribution of the Devonian Domanik facies in the sections of the Russian platform. Sites or boreholes: 1, Plavsk; 2,

Vorotinsk; 3, Pletenevka; 4, Kaluga; 5. Borovsk; 6, Moscow; 7. Povarovka; 8, Redkino; 9. Pachelma; 10, Ukhta; 11,

Bugur.

Legend: 1, limestone; 2, marl; 3, dolomite; 4. anhydrite; 5, gypsum and salt; 6, clay; 7, siltstone; 8. sandstone; 9.

conglomerate; 10, bituminous phtanite; 11, breccian limestone; 12, chert limestone; 13, chert; 14, radiolarian; 15,

unconformity.

FlG. 3.— Distribution des facies Domanik dans les coupes de la Plate-forme russe. Sites ou forages : 1, Plavsk; 2, Vorotinsk; 3.

Pletenevka; 4, Kaluga; 5. Borovsk; 6. Moscow; 7. Povarovka; 8. Redkino; 9, Pachelma; 10, Ukhta; 1 1, Bugur.

Legende : 1, calcaires ; 2, tnarnes ; 3, dolomie ; 4, anhydrite ; 5, gypse et sel ; 6, argile ; 7. pelite ; 8, gres ; 9,

conglomerat; 10, phtanite bitumineuse ; 11, calcaire brechique ; 12, calcaire a cherts ; 13, chert; 14, radiolaires ; 15,

discordance.

The foredeep basins of the Polar Urals and Pay-Khoy are situated in the eastern areas of Timan-

Pechora. The well-known Pechora-Kolva Aulacogen is the axial suture of the Timan-Pechora basin (Fig.

1). It corresponds to a zone of reduced crustal thickness (36-38 km) and high crustal density (up to 2.9

gr/cm 3 ). The basement of the foredeep troughs is depressed up to a depth of 12-14 km (APLONOV &

Shmelev, 1993). The Devonian phase of an active rifting is distinguished in Pripyat trough too

(KLUSHIN, 1992). The Kimmeridgian-Volgian phase is traced to the north into the Barents, Norway and

VALENTINA S. VISHNEVSKAYA

123456789 10

(3 1 2 4b 3 A -$-5 4*6

Fig. 4.— Distribution of the Jurassic domanikoid facies within the sections of the north-eastern part of the Russian platform.

For location of boreholes, see figure 9; for lithology, see legend of figure 3.

Fig. 4 .— Distribution du facies domanikoid jurassique dans les coupes de la partie NE de la plate-forme russe. Pour la

localisations des forages, voir figure 9 : pour la lithologie voir figure 3.

North seas (DYER & COPESTAKE, 1989; Kozlova, 1994) and to the south into the Volga-Urals Basin.

During Devonian time, the spreading rate of the Yapetus was rather high, about 8 cm/year (APLONOV,

1993) which caused a specific type of sedimentation. Rich energy resources are connected with

Devonian and Jurassic rift basins. It is important to emphasize that during rift stages, a thick series of

lacustrine alluvium and active marine marginal slope sediments (Figs. 2-4) locally oil and coal-bearing

have been accumulated. There are numerous examples of lacustrine strata of the Late Devonian

(Ormiston, 1989) and Middle-Late Jurassic age (Fig. 4) which are excellent source rocks. Similar

lacustrine strata have produced oil in the Beatrice field in the North Sea (ORMISTON, 1989). The tectonic

setting in the distribution of the Jurassic domanikoid formations within the Timan-Pechora and Volga

Basins shows resemblance with the same in the North Sea, where tectonic activity along graben

bounding faults were also probably responsible for the character of the deposition of oil and others units

(Dyer & Copestake, 1989).

In the North Sea tectonic rifting caused rapid subsidence which outpaced sedimentation to create

basinal troughs or grabens and the basinal organic rich shales were accumulated in the graben axes

during Kimmeridgian to Ryazanian times (DYER & COPESTAKE, 1989). The characteristic features of

the Timan-Pechora and Volga-Urals rift-basins are also the presence of Domanik and domanikoid-type

bituminous deposits.

DISTINCTIVE FEATURES AND PARTICULARITIES OF THE DOMANIKOID FACIES

Domanik means the Devonian source rocks (oil shales) enriched in organic carbon (kerogen) and

term "domanikoid" is used for others rocks similar to Domanik. The Domanik facies has a special

position in terms of paleogeographic conditions. There are many models for Domanik-type basins. One

of the last was elaborated by ORMISTON (1993). It shows that interest to this problem is high and many

aspects of phenomenon Domanik-type basin formation are still unclear. Oil of the Russian plate is

mostly associated with Devonian Domanik-type rocks (for instance, Pripyat, Timan-Pechora, Volga-

Urals Basins); fuel shale of Peri-Tethys is closely related to Jurassic siliceous bituminous rocks (Timan-

Pechora, Volga-Urals Basins). The primary source of hydrocarbons is Domanik organic shales,

domanikoid-type bituminous limestones and phtanites. The examples of the Ukhta Domanik facies

(Domanik Creek of Ukhta region of Komi, Russia, Figs 2-3, section of site 10) and domanikoid shale of

the Timan-Pechora Basin (Fig. 4, section 4) are given in order to consider this type of specific deposits.

Ukhta Domanik facies (60 m) is represented by grey to black phtanite and siliceous bituminous

Source: MNHN. Paris

DOMANIKOID FACIES OF THE RUSSIAN PLATFORM

limestone with little admixture of clay. The Ukhta middle Frasnian phtanites are oil shale rocks enriched

in organic (kerogen) material. Content of TOC (total organic carbon) approaches 12-20%.

The Jurassic domanikoid-type of sediments is represented by fuel organic shale (Fig. 4, section 4) in

the Volgian stage (up to 1 m in thick) or bituminous clay (Fig. 4, section 9) in the Kimmeridgian (2-3

m). However, the stratigraphic seccession shows that the Jurassic does not have a high TOC (5-10%).

The flooding of former land areas toward the south probably released phosphate and nitrogen into

adjacent waters where they favored the preservation of dead organisms and consequently the

accumulation of TOC due to preservative character of phosphorus (MERTS et al. , 1990).

The chemical and normative composition of Ukhta Domanik facies (MERTS et al., 1990) showed low

Al and alkalies and higher P contents. Mineralogically, the quartz-calcedony prevailed. The secondary

minerals are muscovite-hydromica, montmorillonite, calcite, pyrite, apatite and titanite. The Domanik

oil shale are enriched in Mo, V, Ni. As geochemical indicator of the volcanic material in the phtanites

the montmorillonite has been found. The low content of Mn probably points out the absence of stagnant

environment in the Domanik sedimentary basin and confirms the deep slope condition (MERTS et al.,

1990).

The Jurassic clay of Pechora and Ukhta regions is enriched in glauconite and also montmorillonite.

The glauconite (15-30%), montmorillonite (10-30%), hydromica (10-30%) and chlorite (5-10%) are

dominating in the Kimmeridgian bituminous clay and glauconite (30-40%), kaoline (25-30%),

montmorillonite (20-30%) are prevailing in the Volgian organic rich shale.

Contact of the Domanik facies with the underlying light clay (Fig. 3) is transgressive and begins with

a stratum (about 0.5 m) composed of small pteropod and goniatite shells. The middle most productive

part has a very thin flyschoid structure and in places resembles radiolarites, containing more than 50%

radiolarian skeletons. Upper Devonian (middle Frasnian) sphaeroid-like Entactinia cf. additiva

Foreman, Astroentactinia paronae (Hinde), A cf. crassata Nazarov, A. cf. stellata Naz., A. aff. tantilla

Naz., Haplentactinia inaudita Naz., H. arrhinia For., H. rhinophyusa For., Polyentactinia circumretia

Naz., P. kosistekensis Naz., Tetrentactinia cf. gracilispinosa For., Entactinosphaera grandis Naz., E.

assidera Naz., E. echinata (Hinde), E. nigra (Hinde), E. variacanthina For. (Fig. 5) and spicula-like

forms Paleoscenidium cladophorum Deflandre, Paleothalomnus sp., Ceratoikiscum spinosiarcuthum

For., C. mertsi sp. nov. (Fig. 6) were identified among radiolarians. Domanik radiolarians are associated

with conodonts of Polygnatus timanicum zone. Conodonts Palmatolepis gigas Youngquist, P. subrecta

Miller & Youngquist, P. timanicus Ovnatanova occur in the upper part of domanikoid Sopless suite. The

Palmatolepis punctata conodont zone was determined by A. Yudina (MERTS, 1990) in the middle part

of Domanik Horizon of Ukhta section.

Diversity of radiolaria, isotopic data from brachiopod shells and presence of corals in the

neighbouring areas indicate very warm-water conditions for the Domanik basin. It is in a good

agreement with the latitudinal location of the northern branch of the Yapetus in global reconstructions

near the Equator (APLONOV, 1993). There are a lot of micro-algae Tasmanites in the Domanik facies.

Recent bacteroids and micro-algae occur in abundance within vacuoles in the cytoplasm of living

radiolaria. Older algae may also have been associated with radiolaria and could have existed far from the

coastline for this reason (VISHNEVSKAYA et al., 1993).

The major part of Tasmanites are found in areas located in the neighbourhood of deep faults (Timan-

Pechora, Pyrgidan and Urengoy depressions). Synchronous mass extinction of radiolarians and micro¬

algae can be explained by sudden sharp transgressive-regressive events. The transgressive regime

stimulated an accumulation of Tasmanites because the bulk of this genus was fixed under condition of

transgression maximum (TALNOVA, 1995). For example, the mass deposition of Tasmanites was found

in the delta-front trough of large paleorivers or paleorifts. The thick-walled tasmanitid prasinophyte

algae Tasmanites (ORMISTON, 1989) and radiolarians had an excellent capacity to deliver lipids

(ANDERSON, 1983) for a long time into the bottom sediments. The global affinity of Tasmanites to the

source rocks is very famous (TELNOVA, 1995). On the background of standard methods the co¬

occurrence of Tasmanites and radiolaria can be used for prognosis of oil industry. The absence of

foraminifera is regarded by ORMISTON (1993) as a result of anoxia, similarly to sediments of the

Bonarelli level. The best preservation of radiolarians and algae is probably due to anoxia. The presence

of phosphorus could be served as conservant too.

The upper part of Domanik is cross-bedded and consists of siliceous limestone with abundant micro-

and macrofaunas or polydetritus limestone, and obviously shows a regressive character. The

Kimmeridgian clay (5-45 m) from the Pechora and Ukhta regions has stratigraphic contact with the

VALENTINA S. VISHNEVSKAYA

Fig. 5.— Sphaeroidal radiolariarians from Domanik Horizon

(Domanik Section of Ukhta area. Sample 992).

Radiolaires sphaeroides de l'Horizon Domanik

(coupe du Domanik de la region d'Ukhta, echantillon

992). a, c. e, g, h, i: Entactinosphaera grandis

Nazarov, a, c, e, i x!28: g-h x64; b: E. sp. cf. nigra

(Hinde), x64; d. k: E. echinata (Hinde), d x64; k

x225; f: E. as side ra Nazarov; j: Haplentactinia

inaudita Nazarov, x200.

underlying strata (Fig. 4) and demonstrates the

transgressive-regressive character towards Vol-

gian strata. Radiolarians represent only a minor

part of the total fauna. All taxa are known within

the Boreal province of the Russian platform. The

Kimmeridgian radiolarian assemblage of the

Barents-Pechora-Ukhta region includes Archaeo-

cenosphaera ineaqualis (Rust), Praeconosphaera

ex gr. sphaeroconus (Rust), Pseudocrucella aff.

prava Blome, Crucella crassa (Kozlova), C.

squama (Kozlova), C. aff. mexi-cana Yang,

Orbiculiforma cf. iniqua Blome, O. ? retusa

(Kozlova), Pantanellium tierra-blankaense Pes-

sagno & McLeod, Parvicingula inornata Blome,

P. cf. blowi Pessagno, P. haeckeli (Pantanelli), P.

burnsensis Pessagno & Whalen, P. pizhmica

Kozlova, P. pusilla Kozlova, P. papulata

Kozlova, P. santabarbarensis Pessagno, P. ?

enormis Yang, P. ? blackhornensis Pessagno &

Whalen, Excingula ? bifaria Kozlova.

The study of Parvicingula distribution shows

the predominance of this genus in the Kimme¬

ridgian of the Timan-Pechora and Barents

regions. The species of this genus are represented

by a wide range of morphotypes (Fig. 8). The co¬

occurence of Arcto-Boreal foraminiferal assem¬

blages together with specific group of radio¬

larians and Buchia suggest the possibility to use

Parvicingula as a Jurassic paleoclimatic indicator

(Vishnevskaya, 1996).

The Volgian strata has a transgressive cha¬

racter in the Pechora region. Within middle

Volgian Dorsoplanites panderi ammonite Zone,

among radiolarians Parvicingula papulata

Kozlova, P. conica (Khabakov), P. cristata

Kozlova, P. rugosa Kozlova, P. simplicima

Kozlova are dominating species. Parvicingulids

are up to 90% in the Barents-Pechora region,

along with abundant ammonites, buchias and

foraminiferas in these strata.

The uppermost part of the Volgian has a

regressive character similar to the Ukhta Doma¬

nik Sections. Between Devonian and Jurassic

levels, some similarity exists in the transgressive-

regressive sequences and some difference in the

composition of radiolarian assemblages and their

paleoclimatic affinities. The bulk ot Devonian radiolarians is represented by sphaeroid-like and spicula-

like forms, which could be good deliverers of lipids (VISHNEVSKAYA et al. , 1993), in contrast to the

Jurassic assemblages, where the parvicingulids, probably nassellarians of cold-water upwelling zones,

were dominant. Radiolaria are the only planktic organisms known from Cambrian to recent, and are in

great variety, although of their morphology, evolution, phytogeny and many paleoecological aspects are

still unclear. One question is whether they were planktic, hemiplanktic or benthic during the whole

Paleozoic? As we know, the first attempt of foraminifera to adopt the planktic mode of life started

during Late Paleozoic and Triassic - Early Jurassic times, but appearance of definite planktic forms is

Middle Jurassic (Bajocian) in age.

Source: MNHN. Paris

DOMANIKOID FACIES OF THE RUSSIAN PLATFORM

We do not know if and when radiolarians

passed from benthic to planktic life. Possible

planktic, benthic and transitionals forms could

have existed simultaneously in the Middle-Late

Devonian and possibly up to the Late Jurassic

(Vishnevskaya, 1993). This suggestion has

certain significance for discussion about the time

of appearance of recent oceans and the depth of

the sedimentary domanikoid-type basins.

Bashkirian Domanik facies (30 m) has a

similar structure and differs only by having

phosphatic materials and rare glauconite, which

can indicate a coeval volcanic activity, while

Tatarian one (20 m) is characterized by abundant

brachiopods in its lower and upper parts. To the

north of Ukhta, the Kogva and Kolva Domanik

facies (10-130 m) contain intraformational

sandstone and silstone, indicating possible

influence of islands, and is overlain by bituminous

limestones with lenses of reef limestone debris.

The Jurassic domanikoid facies of Bashkiria

and in the region of Syssola River contain

abundant phosphate. The content of P 2 O 5 reaches

25%, A1 9 0 3 + Fe 2 CK reaches 2-3% and C org.

varies from 2 to 4-5% in the Kimmeridgian clay.

The glauconite represents about 30% of the

Kimmeridgian and Volgian clay. The content of

P 9 0 5 reaches 30%, A1 ? 0 3 and Fe,0 3 is 3-4%,

CaO is 40-45 % and C.o"rg.' is approximately 3-5%

in the Volgian strata. In the region of the Syssola

River in the Volgian Stage, KHUDYAEV (1931)

recognized species belonging to Parvicingula:

(P.) multipora Khudyaev, P. susoUaensis

Khudyaev, P. khabakovi Khudyaev. P. zyrjanica

Khudyaev. The content of parvicingulides in the

Syssola Basin is about 75%.

The Jurassic (Volgian) domanikoid facies of

the Volga-Urals Basin are predominantly

represented by bituminous clay with numerous

fuel organic shale horizons ( 8-10 m, sometimes

15-25 m in thickness). The content of TOC is 10-

55% in the Kashpir Section and 8-22% in the

Gorodische Section. The admixture is composed

of quartz, glauconite, foraminifers, sponge

spicules and radiolarians. Among clay minerals of

the Kashpir Section, hydromica (40-45%) is dominant. The distribution of the clay minerals in the

Volgian organic rich strata of the Gorodische Section is as follows: montmorillonite (20-85%),

hydromica (10-35%), kaolinite (10-45%), glauconite (5-40%). The maximum of montmorillonite is in

the uppermost part of Volgian black shale. The Kimmeridgian clay minerals are represented by

montmorillonite (15-30%), kaolinite (40-45%), hydromica (25-30%) and chlorite-glauconite (5-10%).

In the Gorodische section (Volga Basin), Parvicingula jonesi (Pessagno) is the dominant species

within Kimmeridgian strata and P. blown (Pessagno) is characteristic within Volgian strata. There,

parvicingulid content reaches 50-60%. Within Dorsoplanites panderi ammonite Zone or the uppermost

part of Watznaueria communis nannofossil Zone of the Gorodische Section radiolarians Orbiculiforma

ex. gr. mclaughlini Pessagno, Stichocapsa ? devorata (Rust), Phormocampe favosa Khudyaev,

Parvicingula hexagonata (Heitzer), P. cristata Kozlova, P. conica (Khabakov), P. aff. alata Kozlova,

Fig. 6 .— New species Ceratoiskiscum mertsi sp. nov. (a-e)

and sponge spicules (f-h) from Domanik Section of

Ukhta area. Sample 992. Nouvelle espece

Ceratoiskiscum mertsi sp. nov. (a-e) el spicules

d'eponges (f-h) de la coupe du Domanik dans la

region de Ukhta. a, b, h x225; c, e, f, g x64; d x96.

VALENTINA S. VISHNEVSKAYA

Fig. 7.— Kimmeridgian radiolarians of the Pechora Basin (Ukhta Section, Sample P).

Radiolaires kimmeridgiens du Bassin de Petchora (coupe de Ukhta, echantillon P). a: Archaeocenosphaera ineaqualis

(Rust), xl20; b.e: Praeconosphaera ex gr. sphaeroconus (Rust), b-d, xlOO. e, x200: f: Crucella squama (Kozlova),

x200; g: Pseudocrucella aff. prava Blome. x200: h . j: Pantanellium tierrablankaense Pessagno & McLeod, x250; k:

Crucella aff. mexicana Yang, xlOO; 1: Orbiculiforma sp., xlOO: m: O. cf. iniqua Blome, xlOO; n,o: O. ? retusa

(Kozlova), x100: p: O. sp., xlOO;

P. multipora (Khudyaev), P. aff. haeckeli (Pantanelli), P. aff. spinosa (Grill & Kozur), Plathycrypha-

lus ? pumilus Rust, Lithocampe cf. terniseriata Rust were determined (Figs 9, 10). Tethyan species P.

boesii (Parona) appears in the uppermost part of Volgian stage.

The Volgian strata of the Moscow district is represented by numerous phosphorite or phosphorite

pebbles horizons where the content of TOC is low (less then 5%).

The main Kimmeridgian representatives of the Moscow region are Panncingula vera Pessagno &

Whalen, P. inomata Blome and P. elegans Pessagno & Whalen. Parvicingulids prevails, forming 50%

of this assemblage and in the middle Volgian are represented by P. haeckeli (Pantanelli) and P.

hexagonata (Heitzer) (BRAGIN, submitted).

South-western Saratov Domanikoid facies has organic-rich shale (13 m), while the Pachelma one (to

40 m) is characterized by bituminous clay with Liorhynchus , covered by the Famennian lagune-type

deposits. The western Tambov-Tula Domanik level (40 m) is represented only by bituminous limestone

and clay. To the west, the Devonian strata do not have oil-bearing character, only some coal beds occur.

Kaluga and Moscow domanikoid facies (Fig. 3) are carbonaceous clays (5-38 m) with scarce organic

matter. They are frequently underlain by sulphate complex with reworked spheroid-like radiolarians

indicative of shallow shelf. Pripyat domanikoid facies also differs by numerous shallow-water spongy

radiolarians. Devonian rifting was accompanied by basalt eruptions, as indicated at the Pripyat, Kaluga,

Timan and Volga-Urals sites.

The Domanik basin is interpreted as a relatively shallow-water marginal sea (VISHNEVSKAYA, 1993),

partly separated from Yapetus by an island arc and a barrier reef. Coeval volcanism confirms the active

character of the Domanik basin. The formation of source rocks was concentrated in the Timan-Pechora

and Volga-Urals basins of the Russian Plate, as is observed in the Jurassic rift-type basins. The organic

shales were also accumulating in the same area. For this reasons, the Timan-Pechora and Volga-Urals

Basins are of particular economic interest.

Source: MNHN. Paris

DOMANIKOID FACIES OF THE RUSSIAN PLATFORM

Fig. 8.— Kimmeridgian radiolarians of the Pechora Basin (Ukhta Section. Sample P).

Radiolaires kimmeridgiens du Bassin de Petchora (coupe de Ukhta, echantillon P). a: Parvicingula inomata Blome,

xlOO; b,f: P. aff. bumsensis Pessagno & Whalen, b-c xlOO, d-f, x!20; g: P. haeckeli (Pantanelli), xlOO; h: P. cf. blowi

Pessagno, xlOO; i,k: P. bumsensis Pessagno & Whalen, xl20; 1: P. aff. haeckeli (Pantanelli). xlOO; m.n: P. papulata

Kozlova, xl20; o,p: Excingula ? bifaria Kozlova, o x 150, p xlOO; q.r: E. ? sp.. xlOO; s: Parvicingula ex gr. bumsensis

Pessagno & Whalen, xl20; t.u: P. sp.. xl50; v,w, aa: P. ? enormis Yang, xlOO; x,z: P. ? blackhornensis Pessagno &

Whalen, x, xI50, y-z, x200.

STRATIGRAPHIC CORRELATION OF THE JURASSIC DOMANIKOID FACIES

The lower Kimmeridgian Parvicingula vera Zone of the Barents-Timan-Pechora Basin

(VISHNEVSKAYA & DEWEVER, 1996) is probably equivalent to the lower Kimmeridgian Crucella

crassa Assemblage of KOZLOVA (1994) and correlates with the Buchia concentrica Zone, A. kitchini

ammonite Zone and Epistomina unzhensis foraminiferal Zone as well. This interval may correspond to

Kimmeridgian Clay Hydrocarbon Formation of the North Sea which contains abundant P. jonesi (DYER

& COPESTAKE, 1989). Unfortunately, it is impossible to correlate the details of radiolarian assemblages

and events as well as the stratigraphic distribution of individual taxa, because of confidentiality

involving Britoil and its partners (DYER & COPESTAKE. 1989). For the same reasons, the distribution of

samples within sites or boreholes was not figured.

VALENTINA S. VISHNEVSKAYA

The middle Volgian Parvicingula haeckeli Zone is closely correlated with the Parvicingula papulata

Zone of Pechora Basin (KOZLOVA, 1994) and belongs to the Dorsoplanites panderi Zone which can be

correlated with the Evolutinella emeljanzevi-Trachammina septentrioncilis or Saracenaria pravoslavlevi

foraminiferal Zone (KOZLOVA, 1994) in Pechora Basin and Lenticulina biexcavata Zone (LJUROV,

1994) in Syssola hydrocarbon Basin and Parhabdolithus embergeri nannoplankton Zone in Middle

Volga hydrocarbon Basin. Due to microfossils we can trace this zone in Southern England and Northern

France (VISHNEVSKAYA & DE WEVER, 1996).

Timan-Pechora and Volga Kimmeridgian-Volgian radiolarian assemblages are typically boreal in

character in the presence of Parvicingula and the lack of Mirifusus, Andromeda, Ristola, Acanthocircus

diacranocanthos, but subordinate development of Pantanellidae is fixed (Figs 8, 12). Nassellarian taxa

are dominant in both (Figs 8, 9, 10) assemblages and diversity, mainly represented by species of

Parvicifigula and Stichocapsa.

Fig. 9.— Volgian radiolarians of the Volga-Urals Basin (Gorodische Section, Sample G).

Radiolaires volgiens du Bassin Volga-Oaral (coupe de Gorodische, echantillon G). a,b: Praeconosphaera sp., a x!20, b

xl50 ,c:P ex gr. sphaeroconus (Rust), x!50; d.f: Orbiculiforma ex gr. mclaughlini Pessagno, xl50; g,h: Stichocapsa ?

devorata (Rust). x!50; i.k: Phormocampe favosa Khudyaev. xl50; I: Parvicingula hexagonata (Heitzer), x!50; m,r: P.

? cristata Kozlova, m x200, n-r x!50; s,t: Stichocapsa sp., x!50; u,v:S. sp. A, x!50; w,x: S. sp. B. x!50;

Source: MNHN. Paris

DOMANIKOID FACIES OF THE RUSSIAN PLATFORM

PALEOGEOGRAPHY OF THE VOLGIAN DOMANIKOID BLACK SHALE

The Jurassic stratigraphic sequences of the Timan-Pechora Basin (Fig. 4) show a clearly

transgressive depositional system starting with Early-Middle Jurassic sands and deepening upward to the

accumulation ol the higher grade source rocks in the Volgian time. The mass extinction, observed there

as well as in the Gorodische section of the Volga-Urals Basin (about 40 species of ammonites, 20

species ol aucellids, 22 species of benthic and 20 species of planktonic foraminiferas, 10 species of

belemnites, 5-40 taxa of calcareous nannofossils, 20 species of radiolarians and several species of algae

were recognized within Dorsoplanites panderi Zone), possibly resulted from the cumulative effects of

constant transgressive-regressive episodes (Fig. 4). The schematic paleogeographic reconstruction of the

Volgian time suggests that the eastern rim of the shallow coastal sea had favorable environments for oil

and fuel organic source rock accumulation (Fig. 11). As in recent marine environments, the maximal

concentrations of phytoplankton, siliceous plankton and benthos, carbonaceous plankton, nekton and

benthos were found in the water mass fringering the continent. The relative increase and good

preservation in the bottom sediments of proportions of lipid rich organic matter was related to the

preservative character of phosphorus.

In the Tethyan Realm, the genus Parvicingula is relatively rare, whereas more abundant in the Boreal

(Khytsaev, 1931; Sedaeva & Vishnevskaya, 1995) and Australian provinces (Vishnevskaya, 1996)

reaches up to maximum. Owing to these data we can assume the cold water environments for the

Jurassic domanikoid-bearing basin. Besides, the preponderance of parvicingulides can indicate an

upwelling conditions which could exist along offshore.

PALEONTOLOGICAL DESCRIPTION

The description of some species from Ukhta Domanik Horizon and domanikoid shale is proposed.

Order SPUMELLARIA Ehrenberg. 1875

Family XIPHOSTYLIDAE Haeckel. 1881, emend. Pessagno & Yang, 1989

Genus ARCHAEOCENOSPHAERA Pessagno & Yang, 1989

Archaeocenosphaera inaequalis (Rust), 1898

Fig. 7, a

1898. Cenosphaera inaequalis Rust, pi. 26-1, fig. 6.

1994. Archaeocenosphaera inaequalis (Rust), Kozlova pi. 1, figs 1.2,4-7.

RANGE AND OCCURRENCE— Late Jurassic (early Kimmeridgian), Amoeboceras kitchini ammonite

zone, Ukhta section, Pechora River, Komi area, Russia, Sample P.

Family ACAENIOTYLIDAE Yang, 1993

Genus PRAECONOSPHAERA Yang, 1993

Praeconosphaera ex gr. sphaeroconus (Rust), 1898

Fig. 7, b-e; Fig. 9, c

1898. Conosphaera sphaeroconus Rust. p. 13. pi. 4. Fig. 8.

1993. Praeconosphaera sphaeroconus (Rust). Yang, p. 105, pi. 17, figs 2, 6. 12 ,16 .23.

1994. Praeconocaryomma sphaeroconus (Rust). Kozlova, pi. 1. figs 9. 13.

VALENTINA S. VISHNEVSKAYA

RANGE AND OCCURRENCE— Late Jurassic (early Kimmeridgian), Amoeboceras kitchini ammonite

zone. Ukhta section, Pechora River, Komi area, Russia, Sample P and late Volgian, Craspedites subditus

ammonite zone, Gorodische standard Section of the Russian plate, Sample G.

Praeconosphaera sp.

Fig. 9, a-b

RANGE AND OCCURRENCE — Late Volgian, Craspedites subditus ammonite zone, Gorodische

standard Section of the Russian plate, Sample G.

Family PATULIBRACCHIDAE Pessagno, 1971, emend . Baumgartner, 1980

Genus CRUCELLA Pessagno, 1971

Crucella squama (Kozlova), 1971

Fig. 7, f

1971. Hagiastrum squama Kozlova, p. 1176, text-fig. 1:10.

1994. Crucella squama (Kozlova), Kozlova, pi. 2, fig. 3.

Range AND OCCURRENCE. — Late Jurassic (early Kimmeridgian), Amoeboceras ravni ammonite

zone, Ukhta section, Pechora River, Komi area, Russia, Sample P.

Crucella aff. mexicana Yang, 1993

Fig. 7, k

1993. Crucella mexicana Yang, p. 40, pi. 4, figs 10, 11, 14, 16; pi. 5, figs 10, 21.

RANGE AND OCCURRENCE. — Late Jurassic (early Kimmeridgian), Amoeboceras ravni ammonite

zone, Ukhta section, Pechora River, Komi area, Russia, Sample P.

Family HAGIASTRIDAE Riedel, 1971, emend. Baumgartner, 1980

Genus PSEUDOCRUCELLA Baumgartner, 1980

Pseudocrucella aff .prava Blome, 1984

Fig. 7, g

1984. Pseudocrucella prava Blome, p. 352, pi. 3, figs 1-4, 6. 8-17; pi. 4, figs 1-4, 6-10, 12, 14-16, pi. 15, figs 16-17.

Range AND OCCURRENCE — Late Jurassic (early Kimmeridgian), Amoeboceras kitchini ammonite

zone, Ukhta section, Pechora River, Komi area, Russia, Sample P.

Family PANTANELL1DAE Pessagno, 1977, sensu Pessagno & Blome, 1980

Genus PANTANELLIUM Pessagno, 1977, sensu Pessagno & Blome, 1980

Pantanellium tierrablankaense Pessagno & McLeod, 1987

Fig. 7, h-j

1987. Pantanellium tierrablankaense Pessagno & McLeod, p. 24, pi. 6, figs 5, 7, 8, 13-18.

Range AND OCCURRENCE — Late Jurassic (early Kimmeridgian), Amoeboceras kitchini ammonite

zone, Ukhta section, Pechora River, Komi area, Russia, Sample P.

Source: MNHN. Paris

DOMANIKOID FACIES OF THE RUSSIAN PLATFORM

Pantanellium sp. A

Fig. 12, a

Shell spherical without spines, with a mixture of large-sized pentagonal and hexagonal pore frames.

Only three pore frames visible along diameter of shell.

RANGE AND OCCURRENCE— Late Jurassic (early Kimmeridgian), Amoeboceras kitchini ammonite

zone, Ukhta section, Pechora River, Komi area, Russia, Sample P.

Fig. 10.— Volgian radiolarians of the Volga-Urals Basin (Gorodische Section. Sample G).

Radiolaires volgiens du Bassin Volga-Oural (coupe de Gorodische. echantillon G). a: Parvicingula conica (Khabakov),

xl50; b-f, h-j, m, o-p: P. aff. alata Kozlova, b-f, h, o-p xl50, i-j, m xlOO; g: P. multipora (Khudyaev), xl20; k: P. sp..

xl50; 1: P. aff. thomesensis Pessagno, x!50; n: Parvicingula aff. spinosa (Grill & Kozur), xlOO; q: Platycryphalus ?

pwnilus Rust, xl50; r: Lithocampe cf. terniseriata Rust, x200.

Family ORBICULIFORMIDAE Pessagno, 1973

Genus ORBICULIFORMA Pessagno, 1973

Orbiculiforma cf. iniqua Blome, 1984

Fig. 7, m

1984. Orbiculiforma iniqua Blome, p. 353, pi. 5, figs 3, 4, 6, 7. 10. 11. 14-16.

RANGE AND OCCURRENCE — Late Jurassic (early Kimmeridgian), Amoeboceras kitchini ammonite

zone, Ukhta section, Pechora River, Komi area, Russia, Sample P.

VALENTINA S. VISHNEVSKAYA

Orbiculiforma ? retusa (Kozlova), 1994

Fig. 7, n-o

1994. Staurodictya retusa Kozlova, pi. 2, figs 9-10.

Range and OCCURRENCE— Late Jurassic (early Kimmeridgian), Amoeboceras kitchini ammonite

zone, Ukhta section, Pechora River, Komi area, Russia, Sample P.

Orbiculiforma ex gr. mclaughlini Pessagno, 1977

Fig. 9, d-f

1977. Orbiculiforma mclaughlini Pessagno, p. 74, pi. 4, figs 4-7.

1989. Orbiculiforma mclaughlini Pessagno, DYER & COPESTAKE, p. 224, pi. 1, figs 3-4.

RANGE and OCCURRENCE— Late Volgian, Craspedites subdilus ammonite zone, Gorodische

standard Section of the Russian plate. Sample G.

Orbiculiforma sp.

Fig. 7,1, p

RANGE AND OCCURRENCE— Late Jurassic (early Kimmeridgian), Amoeboceras kitchini ammonite

zone, Ukhta section, Pechora River, Komi area, Russia, Sample P.

Family ENTACTINIIDAE Riedel, 1967, emend. Nazarov and Ormiston, 1984

Genus ENTACTINOSPHAERA Foreman, 1963

Entactinosphaera echinata (Hinde), 1899

Fig. 5, d, k

1899. Heliosoma echinatum Hinde, p. 50, pi. 9, figs 1-2.

1955. Xiphosphaera echinatum (Hinde), Bykova, p. 68, pi. 22, figs 4-5.

1963. Entactinosphaera echinata (Hinde), Foreman, p. 279, pi. 3, fig. 10: pi. 4, fig. 12.

1975. Entactinosphaera echinata (Hinde), Nazarov, p. 60, pi. 3, figs 1-3; pi. 4, figs 1-4.

1983. Entactinosphaera echinata (Hinde), Nazarov & Ormiston, p. 458. pi. 1, figs 6-7.

1993. Entactinosphaera echinata (Hinde), AlTCHJSON, p. 115, pi. 5, figs 6, 11, 14; pi. 7, fig. 3.

RANGE AND occurrence.— Late Devonian (middle Frasnian), conodont zone P. punctata ,

Domanik Horizon, Ukhta section, Domanik Creek, Komi area, Russia, so far as known.

Entactinosphaera assidera Nazarov, 1975

Fig. 5, f

1975. Entactinosphaera assidera Nazarov, p. 64, pi. 5, figs 6-7; pi. 6, figs 6-8.

RANGE AND OCCURRENCE. — Late Devonian (middle Frasnian), conodont zone P. punctata ,

Domanik Horizon, Ukhta section, Domanik Creek, Komi area, Russia, so far as known.

Source: MNHN. Paris

DOMANIKOID FACIES OF THE RUSSIAN PLATFORM

Entactinosphaera grandis Nazarov, 1975

Fig. 5, a, c, e, g-i

1975. Entactinosphaera grandis Nazarov, pi. 5, Figs 11-12; pi. 7, figs 1-4.

1988. Entactinosphaera grandis Nazarov, pi. 14, fig. 8.

1983. Entactinosphaera grandis Nazarov. Nazarov & Ormiston. p. 458, pi. 1. figs 4-5. 1993. Entactinosphaera grandis

Nazarov, Aitchison, p. 115, pi. 5, fig. 8.

RANGE and occurrence. — Late Devonian (middle Frasnian), conodont zone P. punctata,

Domanik Horizon, Ukhta section, Domanik Creek, Komi area, Russia, so far as known.

Entactinosphaera sp. cf. nigra (Hinde)

Fig. 5, b

RANGE AND OCCURRENCE.— Late Devonian (middle Frasnian), conodont zone P. punctata ,

Domanik Horizon, Ukhta section, Domanik Creek, Komi area, Russia, so far as known.

Family HAPLENTACTINIIDAE Nazarov, 1980

Genus HAPLENTACTINIA Foreman, 1963

Haplentactinia inaudita Nazarov, 1988

Fig. 5, j

1988. Haplentactinia inaudita Nazarov, pi. 14, fig. 6.

RANGE and OCCURRENCE. — Late Devonian (middle Frasnian), conodont zone P. punctata ,

Domanik Horizon, Ukhta section, Domanik Creek, Komi area, Russia, so far as known.

Order NASSELLARIA Ehrenberg, 1875

Family PARVICINGULIDAE Pessagno, 1977

Genus EXCINGULA Kozlova, 1994

Excingula ? bifaria Kozlova, 1994

Fig. 8, o-p. Fig. 12, e

1994. Excingula bifaria Kozlova, pi. 5, figs 5-710.

RANGE and OCCURRENCE— Late Jurassic (early Kimmeridgian), Amoeboceras kitchini ammonite

zone, Ukhta section, Pechora River, Komi area, Russia, Sample P.

Excingula sp.

Fig. 8, q-r

Range and OCCURRENCE— Late Jurassic (early Kimmeridgian), Amoeboceras kitchini ammonite

zone, Ukhta section, Pechora River, Komi area, Russia, Sample P.

VALENTINA S. VISHNEVSKAYA

Genus PARVICINGULA Pessagno, 1977

Parvicingula aif.alata Kozlova, 1994

Fig. 10, b-f. h-j, m. o-p

1994. Parvicingula alala Kozlova, pi. 8, Figs 1-3. 6.

RANGE AND OCCURRENCE— Late Volgian, Craspedites subditus ammonite zone, Gorodische

standard Section of the Russian plate. Sample G.

Parvicingula antoshkinae sp. nov.

Fig. 12. h

1984. Parvicingula sp. A, BLOME, p. 364. pi. 9, fig. 10.

DIAGNOSIS.— Test elongate, subconical with nine to eleven chambers. Cephahs spherical smooth,

slightly perforated, with thin and short pointed horn. Thorax, abdomen and first postabdomina chamber

in form of discontinuous circumferential ridges, low and round, formed by rows of thick, Parallel nod

or bars. Postabdominal chambers form subcylindrical part of shell with circumferential ridges. Each

chamber with three rows of polygonal pore frames.

COMPARISON.— Parvicingula antoshkinae sp. nov. differs from P. elegans Pessagno & Whalen,

1982 by presence of nodes and worse development circumferential ridges.

MEASUREMENTS.— (8 specimens): in micrometers. High of test is 280-300, width in apical part

(WA) is 80-90, width of the terminal tube is 135-145.

TYPE LOCALITY.— Clay of Pizhma Creek in Pechora River Basin, Sample P, near Siktivkar town,

Komi area, Russia.

ETYMOLOGY.— Named in honour of Dr. A.I. Antoshkina for her contributions to the study ol the

Timan-Pechora faunas.

RANGE and OCCURRENCE— Late Jurassic (early Kimmeridgian), Amoeboceras kitchini ammonite

zone, Ukhta section, Pechora River, Komi area, Russia, Sample P.

Parvicingula ? blackhornensis Pessagno & Whalen, 1982

Fig. 8, x-z

1982. Parvicingula blackhornensis Pessagno & Whalen, p. 137. pi. 10, figs 10-12, pi. 13. Fig. 14.

1984. Panucingula blackhornensis Pessagno & Whalen, BLOME, p. 357, pi. 9, figs 6, 11, 15, 22; pi. 15, Figs 3, 8.

RANGE AND OCCURRENCE— Late Jurassic (early Kimmeridgian), Amoeboceras kitchini ammonite

zone, Ukhta section, Pechora River, Komi area, Russia, Sample P.

Parvicingula cf. blowi Pessagno, 1977

Fig. 8, h

1977. Parvicingula blowi Pessagno, p. 85, pi. 8, figs 11-14.

1984. Parvicingula blowi Pessagno, PESSAGNO et al. , p. 26, pi. 2, figs 14-15.

1989. Parvicingula blowi Pessagno. Dyer & COPESTAKE, p. 227, pi. 2, figs 3-4.

1995. Parvicingula blowi Pessagno, HULL P- 21, pi. 3, figs 6, 18, 22.

Range and OCCURRENCE— Late Jurassic (early Kimmeridgian), Amoeboceras kitchini ammonite

zone, Ukhta section, Pechora River, Komi area, Russia, Sample P.

Source:

DOMANIKOID FACIES OF THE RUSSIAN PLATFORM

Parvicingula burnsensis Pessagno & Whalen, 1982

Fig. 8, i-k

1982. Parvicingula burnsensis Pessagno & Whalen, p. 136, pi.

1992. Parvicingula cf. bursnensis Pessagno & Whalen, Vishnevskaya, p. 27, pi. 1, fig. 15.

1994. Parvicingula burnsensis Pessagno & Whalen, Kozlova, pi. 3, figs 5-6.

1994. Parvicingula burnsensis Pessagno & Whalen, Vishnevskaya, p. 217, figs 14-19.

Range and OCCURRENCE— Late Jurassic (early Kimmeridgian), Amoeboceras kitchini ammonite

zone, Ukhta section, Pechora River, Komi area, Russia, Sample P.

Parvicingula aff. burnsensis Pessagno & Whalen, 1982

Fig. 8, b-f

Range AND OCCURRENCE— Late Jurassic (early Kimmeridgian), Amoeboceras kitchini ammonite

zone, Ukhta section, Pechora River, Komi area, Russia, Sample P.

Parvicingula ex gr. burnsensis Pessagno & Whalen, 1982

Fig. 8, s

RANGE and OCCURRENCE— Late Jurassic (early Kimmeridgian), Amoeboceras kitchini ammonite

zone, Ukhta section, Pechora River, Komi area, Russia, Sample P.

Parvicingula conica (Khabakov)

Fig. 10, a

1994. Parvicingula conica (Khabakov), Kozlova, pi. 8, figs 4, 8.

RANGE and OCCURRENCE— Late Volgian, Craspedites subditus ammonite zone, Gorodische

standard Section of the Russian plate, Sample G.

Parvicingula ? cristata Kozlova, 1994

Fig. 9, m-r

1994. Parvicingula cristata Kozlova, pi. 4, fig. 11.

RANGE and OCCURRENCE.— Late Volgian, Craspedites subditus ammonite zone, Gorodische

standard Section of the Russian plate. Sample G.

Parvicingula ? enormis Yang, 1993

Fig. 8, w-x, aa; Fig. 12, b

1993. Parvicingula (?) enonnis Yang, p. 118, pi. 19, figs 6, 13, 18; pi. 20, figs 5, 6, 15, 22.

RANGE AND OCCURRENCE— Late Jurassic (early Kimmeridgian), Amoeboceras kitchini ammonite

zone, Ukhta section, Pechora River, Komi area, Russia, Sample P.

VALENTINA S. VISHNEVSKAYA

Parvicingula genrietta sp. nov.

Fig. 12, f

DIAGNOSIS. — Test in form of low conical with short thin horn. Cephalis, thorax, abdomen and

postabdominal chambers covered by irregular pores. Subsequent postabdominal chambers in form of

circumferential ridges with three rows of hexagonal pore frames.

COMPARISON.— Parvicingula genrietta Vishnevskaya, n. sp., differs from P. pizhmica Kozlova,

1994 by the lack of circumferential ridges in the initial part of shell.

MEASUREMENTS. — (4 specimens) mean in micrometers. The high is 210-240, width of apical

portion of the test (WA) is 100-110, maximum width of the terminal tube is 170-200.

TYPE locality.— Clay of Pizhma Creek in Pechora River Basin, Sample P, near Siktivkar town,

Komi area, Russia.

ETYMOLOGY.— Named in honour of Dr. Genrietta Kozlova for her pioner contributions to the study

of the Boreal radiolarians.

Range and OCCURRENCE— Late Jurassic (early Kimmeridgian), Amoeboceras kitchini ammonite

zone, Ukhta section. Pechora River, Komi area, Russia, Sample P.

Parvicingula haeckeli (Pantanelli), 1898

Fig. 8, g. Fig. 12, c

1898. Lithocampe haeckeli Pantanelli, 1880, pi. 10, fig. 6.

1885. Lithocampe haeckeli Pantanelli, RUST, pi. 15, fig. 6.

1971. Eucyrtidium haeckeli (Pantanelli), Kozlova p. 1176, text-fig. 1: 17.

1994. Par\>icingula haeckeli (Pantanelli), Kozlova pi. 3, figs 1-2; pi. 4, fig. 2.

Range AND OCCURRENCE— Late Jurassic (early Kimmeridgian), Amoeboceras kitchini ammonite

zone, Ukhta section, Pechora River, Komi area, Russia, Sample P.

Parvicingula aff. haeckeli (Pantanelli), 1898

Fig. 8 ,1

RANGE and OCCURRENCE— Late Jurassic (early Kimmeridgian), Amoeboceras kitchini ammonite

zone, Ukhta section, Pechora River, Komi area, Russia, Sample P.

Parvicingula hexagonata (Heitzer), 1930

Fig. 9,1

1930. Cyrtocalpis hexagonata Heitzer, p. 391, pi. 1.

Range and OCCURRENCE.— Late Volgian, Craspedites subditus ammonite zone, Gorodische

standard Section of the Russian plate, Sample G.

Source: MNHN. Paris

DOM AN i KOI D FACIES OF THE RUSSIAN PLATFORM

Parvicingula inornata Blome, 1984

Fig. 8, a. Fig. 12, g

1984. Parvicingula inornata Blome, p. 360, pi. 9, figs 7, 17, 18: pi. 10, figs 2. 7, 9, 13.

1994. Parvicingula inornata Blome, Kozlova, pi, 4, fig. I.

1994. Parvicingula inornata Blome, Vishnevskaya, p. 217, figs 14-18.

RANGE and occurrence Late Jurassic (early Kimmeridgian), Amoeboceras kilchini ammonite

zone, Ukhta section, Pechora River, Komi area, Russia, Sample P.

Parvicingula multipora (Khudyaev), 1931

Fig. 10 , g

1931. Dictyomitra multipora Khudyaev, p. 23, pi. 1, fig. 52.

A regularly conical outline of shell with "a spine-like outgrowth"(KHUDYAEV, 1931) or an apical

horn upon the cephalis and three rows of hexagonal pores decreasing in size toward cephalis, as well as

presence ot circumferential ridges at chamber joints allow to consider this species within genus

Parvicingula.

RANGE and OCCURRENCE— Late Volgian, Craspedites subditus ammonite zone, Gorodische

standard Section of the Russian plate, Sample G.

Parvicingula papulata Kozlova, 1994

Fig. 8, m-n. Fig. 12, i

1994. Parvicingula papulata Kozlova, pi. 5, figs 5-7, 10.

RANGE AND OCCURRENCE— Late Jurassic (early Kimmeridgian), Amoeboceras kilchini ammonite

zone, Ukhta section, Pechora River, Komi area, Russia, Sample P.

Parvicingula spinosa (Grill & Kozur), 1986

Fig. 10, n

1986. Pseudodictyom it re l la spinosa Grill & Kozur, p. 253, pi. 7, figs 1-3.

The presense of conical shape with an apical horn and primitive arrangment of pores in horizontal

rings rather indicates the belonging to Parvicingula than Pseudodictyomitrella.

RANGE AND OCCURRENCE— Late Volgian, Craspedites subditus ammonite zone, Gorodische

standard Section of the Russian plate. Sample G.

Parvicingula aff. thomesensis Pessagno, 1977

Fig. 10,1

1977. Parvicingula thomesensis Pessagno, p. 49, pi. 8, figs 12, 25, 29.

Range and OCCURRENCE.— Late Jurassic (late Volgian), Craspedites subditus ammonite zone,

Gorodische standard Section of the Russian plate. Sample G.

VALENTINA S. VISHNEVSKAYA

Parvicingula sp. K

Fig. 12, d

Remarks.— This form differs from all described Parvicingula by presence of well visible vertical

and dorsal spines.

RANGE AND OCCURRENCE— Late Jurassic (early Kimmeridgian), Amoeboceras kitchini ammonite

zone, Ukhta section. Pechora River. Komi area, Russia, Sample P.

Parvicingula sp.

Fig. 8, t-u. Fig. 10. k

Range AND OCCURRENCE— Late Jurassic (early Kimmeridgian), Amoeboceras kitchini ammonite

zone, Ukhta section. Pechora River, Komi area, Russia, Sample P and Late Volgian, Craspedites

subditus ammonite zone, Gorodische standard Section of the Russian plate. Sample G.

Family PHORMOCAMPIDAE Haeckel, 1887

Genus PHORMOCAMPE Haeckel. 1887

Phormocampe favosa Khudyaev, 1931

Fig. 9, h-k

1931. Phormocampe favosa Khudyaev, p. 41, pi. 1, fig. 33.

Range and OCCURRENCE.— Late Jurassic (late Volgian), Craspedites subditus ammonite zone,

Gorodische standard Section of the Russian plate. Sample G.

Family LITHOCAMPIDAE Haeckel, 1887

Genus LITHOCAMPE Ehrenberg, 1838

Lithocampe cf. terniseriata Rust, 1885

Fig. 10, r

1885. Lithocampe terniseriata Rust, p. 315, pi. 39, fig. 13.

1931. Lithocampe terniseriata Rust, KHUDYAEV, p. 25, pi. 1, fig. 44.

RANGE AND OCCURRENCE.— Late Jurassic (late Volgian), Craspedites subditus ammonite zone,

Gorodische standard Section of the Russian plate, Sample G.

Genus PLATYCRYPHALUS Rust, 1885

Platycryphalus ? pumilus Rust, 1885

Fig. 10, q

1885. Platycryphalus pumilus Rust, p. 305, pi. 36, fig. 10.

1931. Platycryphalus pumilus Rust, Khudyaev. p. 39, pi. 1, fig. 17.

Range AND OCCURRENCE.— Late Jurassic (late Volgian), Craspedites subditus ammonite zone,

Gorodische standard Section of the Russian plate, Sample G.

Source: MNHN. Paris

DOMANIKOID FACIES OF THE RUSSIAN PLATFORM

Genus STICHOCAPSA Haeckel, 1881

Stichocapsa ? devorata (Rust), 1885. sensu Dyer & Copestake, 1989

Fig. 9, g-h

1885. Stichocapsa devorata Rust, p. 318, pi. 41, figs 7-8.

1931. Stichocapsa aff. devorata Rust, Khudyaev, p. 45, pi. 1, fig. 46.

1989. Stichocapsa devorata Rust, DYER& Copestake, p. 228, pi. 2, figs 9-11.

RANGE and OCCURRENCE.— Late Jurassic (late Volgian), Craspedites subditus ammonite zone,

Gorodische standard Section of the Russian plate. Sample G.

Stichocapsa sp. A

Fig. 9, u-v

RANGE AND OCCURRENCE.— Late Jurassic (late Volgian), Craspedites subditus ammonite zone,

Gorodische standard Section of the Russian plate. Sample G.

Stichocapsa sp. B

Fig. 9, w-x

RANGE and OCCURRENCE.— Late Jurassic (late Volgian), Craspedites subditus ammonite zone,

Gorodische standard Section of the Russian plate. Sample G.

Stichocapsa sp.

Fig. 9, s-t

Range and OCCURRENCE.— Late Jurassic (late Volgian), Craspedites subditus ammonite zone,

Gorodische standard Section of the Russian plate, Sample G.

Order ALBAILLELLARIA Deflandre, 1953 emend. Holdsworth

Family CERATOIKISCIDAE Holdsworth, 1969

Genus CERATOIKISCUM Deflandre, 1953

Ceratoikiscum mertsi sp. nov.

Fig. 6, a-e

DIAGNOSIS.— Test as with family. Triangular frame is made by a-rod, b-rod and intersector spines.

The a-rod has 12 pairs of caveal ribs. All rods are approximately equal in length. They are armed by the

secondary spines in the post-triangle portions. The patagium and lamellar shell are absent.

COMPARISON.— This species can be distinguished from others due to the presence of numerous (12)

caveal ribs.

Measurements.— (3 specimens): mean (and range) in micrometers. The length of rods is 300 (290-

330). The mean length of caveal ribs is 50.

VALENTINA S. VISHNEVSKAYA

Fig. 11.— Paleogeographic map of the Volgian stage.

1, sandstone and siltstone;2. clay; 3, organic shale; 4, coal and organic detritus; 5, phosphate; 6, land; 7, location of

sections from Fig. 4 (number in circle); 8. location of outcrops (letter in circle).

FlG. 11 .— Carte paleogeographique du Volgien.

1. gres et pelites ; 2, argiles ; 3. shale organique ; 4. charbon et debris organique; 5, phosphate ; 6, terres emergees ; 7.

localisations des coupes de la figure 4 (nombres encercles); 8, localisation des affleurements (lettres encerclees).

Source: MNHN . Paris

DOMANIKOID FACIES OF THE RUSSIAN PLATFORM

Fig. 12.— Some Kimmeridgian radiolarians of the Pechora Basin (Ukhta Section, Sample P).

Quelques radiolaires du Kimmeridgien du Bassin de Petchora (coupe de Ukhta, echantillon P). a: Pantanellium sp. A,

x300; b: Parvicingula ? enonnis Yang, x 170; c: P. haeckeli (Pantanelli), xl70; d: P. sp. K, xl70; e: Excingula ? bifaria

Kozlova, x!70; f: Parvicingula genrietta sp. nov., x215; g: P. inornata Blome, xl70: h: P. antoshkina sp. nov., x215; i:

P. papulata Kozlova, x205.

Source: MNHN . Paris

VALENTINA S. VISHNEVSKAYA

TYPE LOCALITY.— Chert nodules in limestone of Domanik Creek (sample n. 922), near Ukhta town,

Komi area, Russia.

ETYMOLOGY.— Named to honour geologist V. Mens, who first discovered this radiolarian locality.

Range AND OCCURRENCE. — Late Devonian (middle Frasnian), conodont zone P. punctata ,

Domanik Horizon, Ukhta section, Domanik Creek, Komi area, Russia, so far as known.

CONCLUSION

The Domanik and domanikoid facies have the following characteristics: a- decompensated

subsidence on the periphery of Russian platform: b- transgressive-regressive regime; c- enriching in

organic matter.

Domanik and domanikoid facies similar in lithological and geochemical composition but differ in

character of organic matter. The deliverers of the Devonian Domanik were probably algae and

radiolarians in contrast to the Jurassic domanikoid facies, where deliverers were represented by

nannofossils as well as other groups of fauna and flora.

The Devonian Domanik was formed at the west periphery of warm-water Yapetus whereas Jurassic

domanikoid facies were deposited along east rim of Russian Sea within the Boreal province.

Diverse siliceous microfossil assemblages have been obtained from the Volgian strata dated by

ammonites and calcareous nannofossils. Radiolarian co-occurrence with calcareous nannofossils are of

great interest. This description represents the first well-dated radiolarian assemblage of this age at such a

high latitude of the Russian platform.

ACKNOWLEDGEMENTS

The author is grateful for constructive comments made by Fabrice CORDEY and an anonymous

reviewer, which stimulated the improving of this paper. This work is part of the Peri-Tethys program. It

was partly financially supported by the Peri-Tethys Grant N° 95-18 and N°95-96/18 and INTAS Grant

N° 94-1099, RFFI-97-05-6555.

REFERENCES

Anderson, O.R., 1983.— Radiolaria . Springer, Berlin: 1-355.

Aplonov, S., 1993.— Determination of the paleo-oceanic crust age on the base of magnetic and paleomagnetic data: new

examples from the deep sedimentary basin of Russian Arctic shelf. In: L. P. Zonenshain Memorial Conference of Plate

Tectonic, Moscow, november 17-20, 1993. Geomar, Kiel: 30-31.

Aplonov, S. & Shmelev, G., 1993.— Geophysical diagnosis of the suture in the Timan-Pechora basement. In: L. P.

Zonenshain Memorial Conference of Plate Tectonic, Moscow, november 17-20, 1993. Geomar, Kiel: 31-32.

Bragin, N.Y., submitted.— Radiolaria from the phosphorites basal horizons of the Volgian Stage in the Moscow region. Revue

de Micropaleontologie.

DYER, R. & Gdpestake, P., 1989.— A review of latest Jurassic to earliest Cretaceous Radiolaria and their biostratigraphic

potential to petroleum exploration in the North Sea. Northwest European Micropaleontologv and Palynology, London:

214-235.

Khudyaev, J., 1931.— On the Radiolaria in phosphates in the region of the Syssola River. Transactions of the Geological and

Prospecting Service of USSR, 46: 1-48.

Klushin, S.V., Levashov, K.I., Ljzanez, M.G. & Pankova ,V.V., 1992.— The Separation of Gaps and Unconformity'

Boundaries on Seismic and Karatag Diagrammes. Gaps. Unconformities, Nonanticlinal Catchs. Nauka; Technics,

Minsk: 5-11.

Kostiuchenko, S.L., 1993.— The continental plate tectonic model of the Timan-Pechora province based on integrated deep

geophysical study. In: L. P. Zonenshain Memorial Conference of Plate Tectonic, Moscow, november 17-20, 1993.

Geomar, Kiel: 85-86.

Kozlova, G.E., 1994.— Mesozoic radiolarian assemblage of the Timan-Pechora oil field. Proceeding of Saint-Petersburg

International Conference, Saint-Petersburg: 60-75.

Ljurov, S.B., 1994.— Division of Jurassic sequences of the Middle Vichegda. Abstracts of International Symposium

Biostratigraphy of Oil Fields, Saint-Petersburg : 59-60.

Source: MNHN. Paris

DOMANIKOID FACIES OF THE RUSSIAN PLATFORM

* SlE ° m ' V L - 199 °- ° n lhe chemistry of middle Frasnian Ukh.a

Nikishin A.M Milanovsky, E.E., Ziegler, P.A., Lobkovsky, L.I., Cloetingh, S. & Fokin, P.A, 1993.— Devonian rifting

along the eastern and south-eastern margins of European paleocontinent. In: L. P. Zonenshain Memorial Conference of

Plate Tectonic, Moscow, november 17-20, 1993. Geomar, Kiel: 110.

ORMISTON, A.R., 1989. Factors Controlling the Deposition of Late Devonian to Early Carboniferous Oil Source Rocks: a

Global Analysis. Amoco Research Center, Tulsa (USA): 1-56.

ORMISTON, A.R., 1993. The association of radiolarians with hydrocarbon source rocks. Micropaleontology Special

Publication, New-York, 6: 9-16.

Sedaeva, G.M. & Vishnevskaya, V.S., 1995.— Jurassic paleoenvironments of the North-Eastern European platform. In: L. P.

Zonenshain Memorial Conference of Plate Tectonic, Moscow, november 22-25, 1995. Geomar, Kiel: 205.

Talnova, O.P., 1995. Marine phytoplankton from the Devonian rocks of the Timan-Pechora province. In: Ecostratigraphy

and fossil assemblages of Paleozoic and Mesozoic from North-Eastern Europa. Trudy of Komi Scientific Center ,

Siktivkar, 86: 21-30. J

Vishnevskaya, V.S., 1993.— Model of sedimentary basin for the Domanik-type deposits: new version. In: L. P. Zonenshain

Memorial Conference of Plate Tectonic. Vol. 86. Geomar, Kiel: 151-152.

Vishnevskaya, V.S., 1996. Parvicingula as indicator of Jurassic to Early Cretaceous paleogeographical and

sedimentological paleoenvironments within North Peri-Tethys. Abstracts of Moscow Peri-Tethys Workshop, Moscow:

Vishnevskaya, V.S. & De Wever. P., 1996.— About possibility to correlate North Peri-Tethyan radiolarian events with

others zonations. Abstracts of Moscow Peri-Tethys Workshop , Moscow: 31-32.

Paleomagnetism of Permian to Jurassic formations

from the Turan Plate

Marie M. LEMAIRE"', Evgueni L. GUREVITCH ,2 \

Khodjamourad Nazarov 131 , Michel WESTPHAL"',

Hugues FEINBERG 141 & Jean-Pierre POZZI 141

Ecole et Observatoire des Sciences de la Terre, UMR 7516, 5 rue R. Descartes

67084 Strasbourg Cedex, France

<2 ‘ V.N.I.G.R.I., Liteiny 39, 191104 St Petersburg, Russia

,l> Institute of Geology, Academy of Sciences, Ashkhabad Turkmenistan

41 Ecole Normale Superieure, URA 1316, Departement de Geologie, 24 rue Lhomond

75231 Paris Cedex 05, France

ABSTRACT

Paleomagnetic studies have been started upon Upper Triassic-Early Jurassic volcanics and Permo-Triassic redbeds from the

southwest of Turkmenistan. Application of the K/Ar method of dating on volcanic rocks from Turkmenbasi (40°N. 53°E) gives

an homogeneous age of 200 Ma for 5 different lava flows and 227 Ma for an other one. These rocks, distributed among the two

peninsula of Ufra and Turkmenbasi, have a complex magnetization, certainly related to the impressive tectonic history of the

area. We identified a low-blocking temperature component close to the present field, which is probably a recent viscous or

chemical magnetization. The high-blocking temperature component has a more complicated meaning. Declinations may be

affected by possible rotations between sites, and we assume that the metamorphism affected differently the two peninsulas of

Turmenbasi and Ufra. It follows from this that a synfolding component is identifiable among 8 sites from the peninsula of

Turkmenbasi. It has been acquired after a folding of 70%: 1=57° (±11°); pal.=37.6° (± 6 °). This inclination is in a good

agreement with a possible remagnetization during the collision of Iran with the Turan plate in Jurassic time. A low inclination

component, prefolding, with a high-blocking temperature, is identified on the peninsula of Ufra: 1=28.9° (±3.5°); plat.= 15.4°

(±2°). It suggests a shortening of 13° at least, since the Upper Trias. Results from older rocks, Permian (Upper?) redbeds from

Kizil Kaya (4I°N, 55°E), suggest also the importance of the shortening. The high blocking temperature component isolated on

these rocks gives us a reversed direction: D =2 30 ° I=-18°k = 13 a 9 5 = 6 ° ; p 1 a t. =±9° (±3°); VG P: lat. = 34°N

Iong. = l 66 °E (no fold test). Taking into account Permian-Triassic results from Mangyshlak (Kazakhstan), it is highly

probable that a major shortening zone exists between the Turan plate and the stable Eurasian plate north of Caspian Sea.

RESUME

Paleomagnetisme des formations permiennes a jurassiques de la Plaque de Turan.

Letude paldomagnetique de series du Permien-Trias et du Trias superieur-Jurassique inferieur du sud ouest du

Turkmenistan laisse supposer que d’importants raccourcissements ont eu lieu entre la plaque de Turan et l'Eurasie depuis le

Lemaire, M.M., Gurevitch, E.L., Nazarov, K., Westphal, M., Feinberg, H.. & Pozzi J.P., 1998.— Paleomagnetism of

Permian to Jurassic formations from the Turan Plate. In: S. Crasquin-Soleau & E. Barrier (eds), Peri-Tethys Memoir 3:

stratigraphy and evolution of Peri-Tethyan platforms. Mem. Mus. natn. Hist, nat., Ill : 71-87. Paris ISBN : 2-85653-512-7.

Source: MNHN. Paris

MARIE M. LEM AIRE ET AL.

Permien. Deux localites differentes ont ete echantillonnees. La premiere se situe dans la region de Turkmenbasi (40°N, 53°E),

en bordure de la mer Caspienne. a proximite du systeme de failles du Kopet Dagh. Les sites echantillonnes sont distribuSs sur

deux peninsules, Ufra et Turkmenbasi, et sont composes de roches volcaniques calco-alcalines. Ces roches. dont l'age a ete

attribue a 200-227 Ma par la methode K/Ar, portent des directions d'aimantations difficiles h interpreter. Si la composante de

basse temperature de blocage montrant une direction proche du champ actuel peut etre interpretee comme une aimantation

recente visqueuse ou chimique, la composante de haute temperature de blocage est plus complexe. Cependant nous pensons

avoir identifie une composante syn-plissement sur 8 sites de la peninsule de Turkmenbasi. Cette composante aurait ete acquise

apres un plissement de 70%: 1=57° (±11°). La paleolatitude correspondante de 37.6 (±6°) indique que cette re-aimantation

pourrait s'etre produite au cours de la collision de la plaque Iranienne avec la plaque de Turan au Jurassique. Une autre

composante, montrant une faible inclinaison, a ete isolee sur 5 sites de la peninsule d'Ufra: 1=28.9° (±3.5°); pal.=15.4° (±2°).

Cette composante, acquise avant le plissement ou bien apres un faible plissement, suggere un raccourcissement de 13° au

moins, depuis le Trias superieur. La deuxieme localite echantillonnee se situe dans la region du Tuarkyr pres du village de

Kizil-Kaya (41°N, 55°E). Les sites sont representes par des gres rouges pcrmo-triassiques. Une composante de haute

temperature de blocage a ete identiftee dans un site compose de 48 echantillons du Permien superieur: D=230° I=-18° k=13

395=6°; plat.=±9° (±3°); VGP: lat.=34°N long.= 166°E. Cette direction correspond a un raccourcissement de 16° (±6°) depuis la

fin du Permien. Si nous considerons les resultats permo-triassiques de Mangyshlak (FEINBERG et al ., 1996), les resultats du

Permien superieur de Kizil Kaya, et les resultats du Trias superieur de Turkmenbasi, alors il semblerait qu'une zone de

raccourcissement soit presente au nord de la mer Caspienne entre la plaque de Turan et l'Eurasie.

INTRODUCTION

During Permian and Triassic times, the ocean formed between the Scythian-Turanian and Arabian

plates during the last stage of the Pangea assembly, was affected by a major geodynamic evolution.

Microplates, sometimes called "transit plates" (RlCOU, 1995) left the northern margin of Gondwana and

travelled northward, crossing the Paleo-Tethys. Subduction zones under the southern margin of Eurasia

allowed for the disappearance of the former oceanic crust. Sea floor spreading to the south of these

transit plates created the Neotethys. The so-called transit plates collided with the southern margin of

Eurasia during the Late Triassic and Jurassic and were thus generally accounted as responsible for the

Kimmerian phase. Nevertheless the regions east and west of the Caspian Sea were considered during a

long time to be stable parts of the Eurasian plate.

Recently, FEINBERG et al. (1996) and WESTPHAL (1994) studied Permian-Triassic red beds from

Mangyshlak (western Kazakhstan) and showed a large discrepancy between the measured paleolatitude

(D= 12°, 1=31° a 9 5=7°; plat.= 17°(±4°); VGP: 31°N, 206°E) and the expected one for stable Eurasia.

The Mangyshlak area appeared to have moved northward about 10° since Lower Triassic. To obtain

more accurate constraints on this crustal shortening we have undertaken paleomagnetic and

magnetostratigraphic studies farther to the south, in Turkmenistan. Sampling was performed in Permian-

Triassic redbeds from the Tuarkyr erosional window and in volcanics near Turkmenbasi city (formerly

Krasnovodsk).

GEOLOGICAL SETTING

The general setting of the studied area corresponds to the apposition of northern and southern blocks,

separated by a suture. The northern Variscan domain corresponds to the Turan plate, roughly located

north of the Alborz and Kopet Dagh fold belts, and East of Caspian sea. The basement comprises

mainly Lower Paleozoic to Upper Devonian sediments, folded and generally metamorphosed, intruded

by granites dated from 360-370 Ma (ZONENSHAIN, 1990). On the adjoining Scythian plate, the

occurrence of transgressive Carboniferous coal bearing deposits is noticeable. In contrast, the Southern

Gondwanan domain, which corresponds to the Iran and Afghan blocks, is characterized by the complete

lack of Devonian-Carboniferous folding (STAMPFLI, 1978; BERBERIAN & KING, 1981). The

Infracambrian sedimentation is marked by evaporitic deposits, also known on the Arabic plate and in

the Salt Range of Pakistan. Higher, Paleozoic sediments comprise mainly shallow water marine

deposits, with scarce volcanics related to extensional tectonics. The similarity of Lower Devonian

brachiopod faunas from the Arabic platform and from Bandar Abbas (Iran) demonstrates the close

relationships between these two regions, which clearly belong to the same part of the North

Gondawanan margin at this period (SAIDI, 1995). The suture separating the two domains is well

documented in Iran from the west of the Caspian sea (Rasht area) to Aghdarband near Mashad with

numerous remnants of obducted Paleozoic sea-floor (STOCKLIN, 1974; ALA VI, 1991; BAUD et al n

1991). Moreover, chemical analysis of ophiolites from Talesh and Mashad reveals that they were

PALEOMAGNETISM OF PERMIAN TO JURASSIC FROM TURAN PLATE

RUSSIA

study a

TJRK/v,

KAZAKHSTAN

KARA-BOGAZ

CASPIAN

GULF

TUARKYR

Turkmenbasi

Permian

Triassic

Jurassic

□ Cretaceous □ Quaternary

I:;;:;; Paleogene ★

n Neogene

CS Faults, according

to Erteleva et al. (1994)

Triassic Jurassic

volcanism

Fig. 1. — Geological outline of the studied area, according to Krimous et al. (1989) and Erteleva^/ al. (1994)

Fig. 1 .— Carte geologique de la region etudiee, d'apres KRiMOUSei al. (1989) et Erteleva et al. (1994).

MARIE M. LEMAIRE ET AL.

originally close to a subduction zone (WEBER-DiEFENBACH et al, 1986). This suture was also

recognized in northern Afghanistan in the Paropamisus mountains by BOULIN & BOUYX (1977 a,b).

After the Palaeozoic evolution described above, the Turan plate was affected during Late Permian

(?) and Triassic times by a rapid subsidence, with the sedimentation of a thick molasse. At about the

same time, the Southern margin of the plate experienced an important volcanic activity, with lava and

pyroclastic deposits evolving from acidic (rhyolite-dacite) to andesitic type. This intra-plate volcanism

is related to the subduction of Paleo-Tethys before and just after the ‘docking’ of the Iran block. On the

Turan plate the first Kimmerian folding occured during the Late Triassic or Early Jurassic, and

continued throughout Jurassic times (FEINBERG et al., 1996). Recently, during the Pliocene tectonic

phase, the southern part of the Turan plate has been truncated by a major fracture called the Kopet-

Dagh-Balkan fault (Fig. 1), which corresponds to the overthrusting of the Kopet-Dagh structures on the

Turan plate (ARTEMJEV & Kaban, 1994). This active fault, oriented N120° (LYBERIS et al, 1995),

plays an important seismic role with high magnitudes earthquakes (Krasnovodsk, 1895; Ashkhabad,

1948). The Kopet-Dagh fault has been sometimes presented as the southern boundary of the Turan

plate, but this fault is in fact located 100 to 300 kilometers north of the northern border of the

Gondwanan ‘mega-terrane' which is the true limit.

SAMPLING

Sampling was carried out in two regions just north of the Kopet Dagh which have Mesozoic and

Paleozoic outcrops (Fig. 1). The first region of Turkmenbasi and Ufra peninsulas, shows Triassic-

Jurassic volcanics unconformably overlain by Jurassic sediments. Each site represents one or several

units: a lava flow, a pyroclastic deposit, or a volcanoclastic deposit. The samples are distributed over

several tens of meter of outcrop in order to avoid local tectonic or magnetic perturbations. Between 5 to

15 samples were cored with a portable drill or hand sampled. Orientation was done with a magnetic

compass and the coordinates of the sites determined with a G.P.S. (Global Positioning System) receiver.

After paleomagnetic analysis, some very close sites were grouped together into single statistical units.

Bedding planes, fluidal structures were measured whenever it was possible. Separate hand samples of

volcanics were also taken for age determination and geochemistry.

The second region, was near Kizil Kaya, in the Tuarkyr region (Fig. 1), where Permo-Triassic

redbeds outcrop there. Sampling was organized into long stratigraphic section which covers the

Amanbulak series (Permian) and the beginning of the Lower Triassic. The Amanbulak section has a

thickness of 3 km at least. It is composed of sandstones, clays and conglomerates.

GEOCHEMISTRY

Six different lava flows from the Turkmenbasi area, were sampled in order to determine their

radiometric ages with the K/Ar method. We .selected, within the largest geographic scattering from Ufra

and Turkmenbasi peninsulas, a representative fan of the various petrographic types. Chemical

classification and nomenclature of the samples has been established on the basis of Na 2 0+K 2 0 content,

relative to the content of Si0 2 (Le Bas et al ., 1991). The acidic rocks are representative of a

calcalkaline volcanism (BROUSSE, 1971; GlROD, 1978). Results of major- and trace-element analyses

give a composition of the magmas similar to those which erupt at convergent plate boundaries (GILL,

1981). The sites selected have also significant paleomagnetic results. Two very close samples (UEA03

and UEB08) were chosen because they show opposite polarities.

Whole-rock analyses were obtained from the 100 to 160 pm sieve fraction. Potassium was measured

by flame photometry with a lithium internal standard. Argon was extracted in a bakeable glass vacuum

apparatus and determined by isotope dilution techniques (using 38 Ar as a tracer) in a MS 20 mass

spectrometer. All samples were treated by the static method. The set constants recommended by

Steiger & Jaeger (1977) were used for age calculation.

K-Ar results are presented in Table 1. Five samples show an homogenous age around 200Ma. Site

TWA is older with an age of 226.8 Ma. Supplementary datations using 39 Ar/ 40 Ar methods are in

process in order to control if this divergence is significant; wether it is real or related to the method of

the K/Ar dating.

Source: MNHN, Pans

PALEOMAGNETiSM OF PERMIAN TO JURASSIC FROM TURAN PLATE

Table I — lOAr dating of volcanic rocks from Turkmenbasi and Ufra (whole-rock). The nomenclature is based on theLEBAS

etal. (1991) classification, wt.%: percent of weight.

Tableau /.— Duration par la methode K/Ar des roches volcaniques des peninsules de Turkmenbasi el Ufra. La nomenclature

des roches esl basee sur la classification de LF.BASet al. (1991). wt.% : pourcentage du poids.

Samples

or sites

Location

Petrographic

type

K 2 0

(wt.%)

IQ0rad 4 °Ar

Tolal 40 Ar

rad.^Ar

(10' 1 hnol/g)

Age(±a)

in Ma

UEA03

39°59 , 48"N-53 o 06 , 22 , 'E

trachyte

2.936

96.9

86.44

I93.7±4.4

UEB08

39°59'48"N-53 o 06'22”E

dacite-rhyolite

3.543

95.8

110.74

205.0±3.0

UWB 06-07

39°58 , 27”N-53 o 03'27"E

rhyolite

3.482

97.1

103.36

195.2±2.9

TWA

40°01' 18"N-52 0 55'44 0 E

dacite

2.222

96.6

77.32

226.8±5.2

TEA

40 o 00'46"N-52°56'59"E

dacitc

2.936

976

87.21

195.3±4.5

TCA

40 o 00 , 28°N-52°57 , 04"E

rhyolite

4.535

97.9

141.43

204.6±3.0

PALEOMAGNETISM

Procedures and measurements

The drilled cores and hand samples were cut to the standard size of paleomagnetic specimens and

measured in the Strasbourg and Paris paleomagnetic laboratories with a CTF cryogenic magnetometer

or a modified Digico spinner with an equivalent noise level of 3.I0- 5 A/m. Susceptibilities were

measured with a Digico bulk susceptibility bridge, with a sensitivity of 2 10' 6 (S.I.).

We tested both alternating field and thermal demagnetization methods on pilot specimens. With the

A.F method, parasitic magnetizations are often observed aboye 80mT and hamper the determination of

the final magnetization. Therefore the thermal method was preferred for the main part of the samples

Fourteen steps of heating were made up to 590-600°C for Turkmenbasi, or up to 620°C for Kizil-Kaya

Susceptibilities were controlled every two steps to detect any mineral transformation which could alter

the final direction. Demagnetization curves were analysed by principal component analysis on

stereonets and on Zijderveld diagrams (ZlJDERVELD, 1967; KlRSCHVlNK, 1980; KENT et al., 1983).

Fisher statistics (FISHER, 1953) were used to compute the average direction of each site and its

statistical properties.

Within-site fold tests (McElhinny. 1964; Watson & Bjkin, 1993) were usually not possible due

to lack of attitude variation of strata within each site. However between-site fold tests were possible. A

reversal test was possible in one volcanic site from Ufra.

Results

Turkmenbasi area

Specimens distributed among 19 sites were demagnetized. Different behaviour was observed from

one site to another. The NRM intensities vary globally from 10‘ 3 A/m to lOA/m. In each site this

intensity varies only by a factor 1 to 20. The Koenigsberger factor was calculated in order to identify

any anomalous specimen; it varies from 0.1 to 4.

During thermal demagnetization one or two components of magnetization can be distinguished (Fig.

2). A low temperature component destroyed below 350°C is isolated at 5 sites. The average directions

are reported in Table 2. They are close to the present field direction, which is: D=5° 1=59°. Among the

other 14 sites, the low component temperature, which sometimes appears, is too scattered to be

analysed.

MARIE M. LEM AIRE ET AL.

Table 2.— Low-blocking temperature components. Dg, declination in geographic coordinates; Ig, inclination in stratigraphic

coordinates; kg, precision parameter in geographic coordinates; 395 , semi-angle of the cone of 95% confidence in

degree.

Tableau 2 .— Composantes de basse temperature de blocage. Dg, declination en coordonnees geographiques ; Ig, inclination

en coordonnees geographiques ; 095 demi-angle du cone de confiance a 95%, en degres.

SITE

number

of samples

polarity

blocking

temperature

•g

TURKMENBASI

TWC

200

TWE

200-250

336

344

TEAB

200-250

345

UFRA

<350

KIZIL KAYA

<350

355

<350

355

<350

MEAN

number

of sites: 9

357

5 9

A high temperature component destroyed above 500°C is quite well defined for 13 sites (Table 3,

Fig. 3, Fig. 4). It is carried by magnetite. Because of heterogeneous behaviour of some samples and

scattered directions, the characteristic mean directions of 6 sites could not be isolated.

Twelve of the well defined final directions are normal. Site UEAB shows both normal and reversed

directions (Fig. 5). A reversal test is possible on them. The antipodal angle lies between 160 and 170°

(Table 4) and, according to McFADDEN & McELHINNY (1990), classifies this lest as C. The test is

positive, but the slight deviation indicates that a secondary contamination is still present in the primary

magnetization. Nevertheless we are not in presence of a massive remagnetization.

We performed then a progressive fold test (WATSON & ENK1N, 1993) on mean directions from Ufra

and Turkmenbasi (Fig. 6) all together. The best grouping is found for uncorrected or slightly corrected

directions (Table 5). The k and a 95 are not very satisfactory, but at this point of our analysis we could

suppose that this mean is a secondary magnetization, and that the scattering is related to local rotations.

Then, to avoid the contribution of possible rotations, we performed the statistical analysis of

inclination data only of ENKIN & WATSON (1996) (Fig. 7a). A different conclusion appears (Table 6).

The strongest scattering is now in geographical coordinates. But it is meaningless to conclude from then

on that the direction is primary. This progressive fold test is not conclusive.

Should we take into account the best grouping firstly found (Table 5), and then consider that all the

samples have been remagnetized? The presence of a non negative reversal test and of an insignificant

fold test on inclinations only, let us suppose that things are not so simple. The scattering could be

related to an heterogeneous remagnetizing contribution.

We notice that the 5 sites from Ufra peninsula, representing independant Hows, have a common

behaviour; they all show inclinations lower than 40°, both before and after tectonic corrections (Fig. 4).

The sites from Ufra are separated from Turkmenbasi ones by 7 km at least. For this reason, we decided

to separate these two peninsula, supposing that metamorphism has affected differently these two

localizations.

Source: MNHN, Paris

PALEOMAGNETISM OF PERMIAN TO JURASSIC FROM TURAN PLATE

N, up

Fig. 2.— Examples of demagnetization curves, o, horizontal plane; +, East-West vertical plane. Characteristic temperatures are

indicated.

Fig. 2.— Exemples de courbes de desaimantation. o, plan horizontal ; +, plan vertical Est-Ouest. Les temperatures

caracteristiques sont indiquees.

Source: MNHN. Paris

MARIE M. LEMAIRE ETAL.

before

tectonic

correction

/ partial

f olding

(70%) i

after

tectonic

correction

Fig. 4.— Distribution of the high blocking temperature component for

5 sites from the peninsula of Ufra before tectonic correction (a)

and after tectonic correction (b). The grey circle indicates the

mean inclination after the tectonic correction.

Fig. 4 .— Distribution de la composante de haute tempertaure pour les

5 sites de la peninsule d'Ufra avec la correction tectonique (a) et

apres la correction tectonique (b). Le cercle gris indique

I'inclinaison moyenne pour apres correction de pendage.

Fig. 3.— Distribution of the high blocking temperature component for

8 sites from the peninsula of Turkmenbasi, (a) before tectonic

correction; (b) after a 70% folding; (c) after a complete tectonic

correction. The grey circle indicates the mean inclination for a

partial folding of 70%.

F/G. 3 .— Distribution de la composante de haute temperature pour les

8 sites de la peninsule de Turkmenbasi, (a) avant la correction

tectonique ; (b) apres un plissement partiel de 70% (c) apres la

correction tectonique. Le cercle gris indique I'inclinaison

moyenne pour un plissement partiel de 70%.

Source: MNHN. Paris

PALEOMAGNETISM OF PERMIAN TO JURASSIC FROM TURAN PLATE

Table 3.— High blocking temperature components. Dg/s, declination in geographic/stratigraphic coordinates; Ig/s, inclination

m ge °f r r?c P ^ ,c/ s^ngraphic coordinates; ks, precision parameter in stratigraphic coordinates; aos semi-angle of the

cone ot 95% confidence, in degree, in stratigraphic coordinates; s, strike; d, dip.

Tableau 3.— Composantes de haute temperature de blocage. Dg/s, declinaison en coordonnees geographiaues /

stratigraphiques; Ig/s, inclinaison en coordonnees geographiques / stratigraphiques ; a 95 demi-angle du cone de

confiance a 95% en coordonnees stratigraphiques ; s, strike; d, dip.

SITE

position

number

of samples

polarity

blocking

temperature

cx95

TURKMENBASI

TWA

40°0L3 , N-52°55.7 , E

580

144

115

TWC

40°01.2‘N-52°55.9'E

570-580 80

166

115

TWE

40 o 0fN-52 o 55.9E

585

164

109

40°00.7'N-52°57E

580-60075

100

202

40°00.7'N-52°57.1 'E

590

265

40°00.7’N-52°57.8E

600

133

103

TEAB

40°00.8'N-52°57E

585

344

TCA

40 o 00.5'N-52°57.1 'E

590

329

115

UFRA

UWA

39°58.4'N-53°03.3 , E

570-585 76

UWB

39°58.5 , N-53°03.3’E

570-585 79

102

UOB

39°59.05'N-53°03.5E

600

UCD

39°59’N-53°03.6E

600

104*

114

UEAB

39°59.8 , N-53°06.4 , E

35%R

500-520 39

KIZIL

KAYA

40.5°N-55.5°E

92%R

230

-18

150

Table 4.— Reversal test on site UEAB. We use the McFadden & McElhinny (1990) classification. For explanation of

symbols see Table 2.

Tableau 4.— Test des inversions sur le site UEAB. Nous utilisons la classification de McFadden & McElh/NNY (1990). Voir le

tableau 2 pour Vexplication des symboles.

Stratigraphic

coordinates

Number

of samples

Declination

Inclination

«95

Reversed polarity

230

-30

Normal polarity

29.5

both

Angle between normal and reversed directions

(yo)

7.4°

Critical angle

(ye)

18°

Test classified Rc

Source

MARIE M. LEMAIRE ETAL.

Fig. 5.— Distribution of the antipodal polarities of the samples

from site UEAB.

FlG. 5. —Distribution des directions antipodales du site UEAB.

Fig. 6.— Watson & Enkin (1993) fold test on Turkmenbasi

and Ufra peninsulas. Horizontal: unfolding ratio. The

curves show 20 examples of variation of k (precision

parameter). The histogram gives the position of the

best unfolding ratio.

Fig. 6 . —Test du pli de Watson & FNkin (1993) applique aux

peninsules de Turkmenbasi et d'Ufra.

Horizontalement : taux de deplissement. Les courbes

representent 20 exemples de variations de k,

parametre de precision. L'histogramme donne la

position du meilleur taux de deplissement.

Table 5.— Watson & Bjkin (1993) fold test on Turkmenbasi and Ufra peninsulas. For explanation of symbols see Table 2.

Tableau 5 .— Test du pli de Watson & Enkin (1993), applique aux peninsules de Turkmenbasi et Ufra. Voir le Tableau 2 pour

l'explication des symboles.

Number

of sites 1 3

Declination

Inclination

a y5

in situ

65.4

46.4

10.6

12.8

complete unfolding

117.5

46.9

5.4

17.9

partial unfolding.

4% (±10%)

67.1

47.3

10.6

12.8

Source: MNHN. Paris

PALEOMAGNETISM OF PERMIAN TO JURASSIC FROM TURAN PLATE

Fig 7.— a, Enkin & Watson (1996) statistical analysis on

inclination only. Turkmenbasi and Ufra. For

explanation see Fig. 4. b, Enkin & Watson (1996)

statistical analysis on inclination only: Ufra. For

explanation see Fig. 4. c, Enkin & Watson (1996)

statistical analysis on inclination only: Turkmenbasi.

For explanation see Fig. 4.

Fig. 7.—a, Analyse statistique de Enkin & Watson (1996) sur

les inclinaisons seules. Turkmenbasi el Ufra. Pour

Texplication de la figure voir Fig. 4. b. Analyse

statistique de Enkin & Watson (1996) sur les incli¬

naisons seules: Ufra. Pour Vexplication de la figure

voir Fig. 4. c. Analyse statistique de Enkin & WATSON

(1996) sur les inclinaisons seules: Turkmenbasi. Pour

Texplication de la figure voir Fig. 4.

% unfolding

Table 6.— Enkin & Watson (1996) statistical analysis on

inclination only; Ufra and Turkmenbasi peninsulas,

a, standard deviation.

Tableau 6 .— Analyse statistique sur les inclinaisons seules

d'apres Enkin & Watson (1996) ; application aux

peninsules d'Ufra et de Turkmenbasi. s, ecart-type.

Number

of sites 1 3

Inclination

in situ

43.7

17.3

complete unfolding

36.8

15.2

partial unfolding

106% (± 52%)

35.1

15.0

Table 7.—- Enkjn & Watson (1996) statistical analysis on

inclination only; Ufra peninsula, a, standard devia¬

tion.

Tableau 7.— Analyse statistique sur les inclinaisons seules

d'apres Enkin & Watson (1996) ; application a la

peninsule d'Ufra. s, ecart-type.

Number

of sites 5

Inclination

in situ

28.4

7.3

complete unfolding

28.4

3.6

partial unfolding

28.9

3.5

88% (± 36%)

Source: MNHN. Paris

MARIE M. LEMAIRE ETAL.

Table 8 . — Enkin a& Watson (1996) statistical analysis on

inclination only; Turkmenbasi peninsula, a. standard

deviation.

Tableau 8 .— Analyse statistique sur les inclinaisons seules

d'apres Enkin & WATSON (1996); application a la

peninsule de Turkmenbasi. s, icart-type .

Number

of sites 8

Inclination

in situ

53.2

14.4

complete unfolding

42.2

17.5

partial unfolding

57.1

11 0

30% (± 13%)

UFRA PENINSULA.— The distribution of

mean site directions from Ufra shows clearly

that this distribution is not fisherian, neither

before or after tectonic correction. Site

UEAB has a similar inclination but a quite

different declination indicating possible local

tectonics rotations. We performed the Enkin

and Watson test (1996) on the inclinations.

The inclinations keep beeing low when we

applied the tectonic correction. But the

grouping improves (Fig. 7b) significantly for

a partial or complete correction (Table 7).

The mean inclination of 28.9° (±3.5°) could

be either primary or synfolding.

TURKMENBASI peninsula.— The same procedure was done with Turkmenbasi results (Fig. 3). The

best grouping is observed after a partial tectonic correction (Fig. 7c; Table 8 ). The mean inclination of

57.1° (±11°) is a remagnetization because the best grouping is found for a partial folding of 70%

(±13%).

KIZILKAYA

Five sites have been sampled and 120 specimens were demagnetized up to 620°C (Fig. 2). Only

section KC, which is the longest, gave good results. This section represents a thickness of 1km almost,

and is stratigraphicaly situated just below the Permian-Triassic boundary, in the upper part of the

Amanbulak section. The NRM ranges from 1 to 500 mA/m. The first component has an unblocking

temperature lower than 350°C. Its direction, well defined, is close to that of the present field (Table 2).

It is probably a recent viscous or chemical remagnetization due to the weathering of the samples.

Above this temperature one or two components exist. Magnetization is heterogeneous: some samples

see their magnetization disappearing below 580°C and for others 50% of the magnetization remains

above 620°C. Only site KC gives consistent results. The large majority of the samples are of reversed

polarity. A few samples are normal, but they are too few and the scatter is too large for a reversal test.

An internal fold test is not significant because the bedding-plane attitudes are very similar. The

characteristic magnetization direction is (Table 3) in stratigraphic coordinates: D=230° I=-18° k=13

a 9 5 = 6 °; plat.=9°; PGV: lat.=34°N long.= 166°E.

DISCUSSION

Turkmenbasi area

In view of the result of the progressive fold test on inclinations only (Fig. 7a), performed on the 13

final directions of Turkmenbasi and Ufra peninsulas, a pre-folding component is not identifiable. But

this test neither proves that all these final directions have been remagnetized.

We consider another solution which is based on the heterogeneousness of the remagnetization.

However, a detailed petrographic study must confirm the hypothesis of a different metamorphism in the

Turkmenbasi and Ufra volcanic series.

Our reasoning is based on inclinations only. We cannot trust the declinations because of the eventful

tectonic story of this area, close to the Ashkhabad fault (Fig. 1). The Enkin and Watson (1996) statistical

analysis of paleomagnetic inclination data, allows us to isolate two components:

— a low inclination: 28.9° (±3.5°). This low inclination could either be prefolding or contemporary

with the beginning of the folding (Table 7). It corresponds to a paleolatitude of 15.4° (±2°) (n °8 on Fig.

Source: MNHN. Paris

PALEOMAGNETISM OF PERMIAN TO JURASSIC FROM TURAN PLATE

— a synfolding inclination, 57.1° (±11°), which corresponds to a remagnetization at the end of the

folding (folding of 70%). It implies a paleolatitude of 37.6°(±6°) (n°17).

We calculated the theoritical paleolatitudes for a point in the middle of the studied area (Turkmenbasi

city: 41°N, 53°E). For Eurasia and Africa we used the apparent polar wander curve of BECK (1994) and

draw the expected paleolatitudes since 280Ma (Fig. 8). We have to remind that a first folding occured

before the Jurassic because the volcanics series from Turkmenbasi are unconformably overlain by

Jurassic sediments. Therefore, we attribute an age comprised between 200 and 230 Ma, if we consider

that the older age of 227Ma is significant, to the low inclination (n°8). It can be inferred that a

shortening of 13° at least, occured since this the Upper Triassic.

The synfolding component, corresponding to a paleolatitude of 37.6° (±6°), is in a good agreement

with the hypothesis of a remagnetization of the volcanic series at the end of the collision of Iran with the

Turan plate (Fig. 8).

KIZIL KAYA

Section KC, which is thought to be Upper Permian, has a main direction of: D=230° I=-18°; plat.=

±9° (±4°). The expected direction if this Tuarkyr region was part of Eurasia 260Ma ago is: D=45° 1= 45°

(BECK, 1994). Then a paleolatitude of -9° in the southern hemisphere corresponds to a shortenning of

25° (±6°) and to a rotation of 175° (±9°) with respect to Eurasia, in this case the main direction is

reversed (D=230° I=-18°). Whereas a paleolatitude of +9° (±4°) corresponds to a shortenning of 16°

(±6°) and to a rotation of 5° (±9°) only. In this case the main direction is normal (D=50° 1=18°). We

think that the second hypothesis, a paleolatitude in the northern hemisphere, is more likely because the

Tuarkyr region isn't highly affected by tectonic movements: a rotation of 5° (±9°) seems to be more

probable than a rotation of 175° (±9°).

Thus the main reversed direction observed probably represents one of the reversed zones within the

Illawara mixed polarity zone and we assume that a shortening of 16° (±6°) occured since the Upper

Permian between Turan plate and Eurasia.

CONCLUSION

We reported our results and also existing ones from the Iran and Turan plates (Table 9), for the

Permian, Trias and Jurassic, on a diagram of the expected paleolatitudes (Fig. 8). According to SOFFEL

& FORSTER (1980) and WENSINK (1981, 1982), we consider a belonging of Iran to the Gondwana

during Permian and Post Permian times. Then the paleolatitudes for the Alborz (n°l on Fig. 8), Abadeh

& Jamal (n°2) and Sorkh Shales (n°7) results are drawn in the southern hemisphere (Fig. 8). The

majority of the results are significantly below the theoretical curve of Eurasia. From the Permian to the

Jurassic, the paleolatitudes increase progressively to join the curve of Eurasia at the end of the Jurassic.

We gave two results for the Cretaceous from Kopet Dagh and Alborz, which show their belonging to

Eurasia at this time.

Results n°5 and n°13 are distinct from the other. It isn't the aim of this paper to discuss the validity of

the paleomagnetic data, but we have to notice that the previous results from Turkmenistan are

heterogeneous. The paleomagnetic analysis are sometimes made on samples which are not

demagnetized, and without any fold or reversal test.

Excluding results n°5 and n°13, we observe that the difference between results from the Turan plate

and Iran is not large and may be not significant after the Permian. We suppose that a major shortenning

zone between Iran, Turan and the stable Eurasian plate might exist somewhere north of the Caspian Sea.

ACKNOWLEDGEMENTS

Financial support was given by the Peri-Tethys program (94-27), INTAS (94-3036) and CNRS

(MDRI). Geochemical analysis and K-Ar age determinations were made by the Centre de Geochimie de

la Surface (CNRS-UPR 6251 Strasbourg). This is a contribution of CNRS, UMR 7516 and URA 1316.

MARIE M. LEMAIRE ETAL.

Paleolatitude

80 120

CRETACEOUS

16 C

20C

JURASSIC

TRIAS

Europe

Africa

Turkmenistan

Iran

Kazakhstan

(5) Turkemia (M1981)

(b) Kizil-Kaya (E)

j primary component

^j, secondary component

(7) Sorkh Shales (B)

(T5) Turkmenia (M1016

0*7) Turkmenbasi (E)

Fig. 8 .— Theoretical paleolatitudes according to the apparent polar wander curves for Eurasia and Africa (Beck, 1994).

Crosses show the values obtained in different places with errors. References in Table 9 are indicated by number 1 to 17.

For further explanations see Table 9.

Fig. 8 .— Paleolatitudes theoriques d'apres les courbes de derive des poles d'Europe et d'Afrique (BECK 1994). Pour plus

d'informations voir le tableau 9.

Source: MNHN. Paris

PALEOMAGNHTISM OF PERMIAN TO JURASSIC FROM TURAN PLATE

Table 9.— Paleomagnetic results from Iran, Kazakhstan and Turkmenistan for Permian, Trias and Jurassic. The paleolatitudes

were calculated for a point at the reference of Turkmenbasi. (I): Iran; (K): Kazakhstan; (T): Turkmenistan; ?: post¬

folding component; lat.: latitude; long.: longitude; paleolat.: paleolatitude; dpal.: error bar on the paleolatitude; ref.:

reference of the pole; (A): Soffel & Forster , 1980; (B): Wensink, 1979, 1982; (C): Wensink & Varenkamp, 1980;

(D): Feinberg et al., 1996; (E): this paper; (F 7 ): Bazhenov, 1987; M: McElhinny data base.

Tableau 9 — Resultats paleomagnetiques d'lran, du Kazakhstan et da Turkmenistan pour le Permien, le Trias et le Jurassique.

Les paleolatitudes (paleolat.) ont ete calculees pour un point de reference situe a Turkmenbasi. (I) : Iran; (K) :

Kazakhstan ; (T) : Turkmenistan ; ? : composante post-plissement; plat. : marge d'erreur sur la paleolatitude ; ref.

reference des poles ; (A) : SOFFEL & FORSTER, 1980 ; (B) : Wensink, 1979, 1982; (C) : Wensink & Varenkamp, 1980;

(D) : Feinberg et al ., 1996; (E) : cel article; (F) : Bazhenov, 1987; M ; base de donnees de McElhinny.

REGION <

Coordinates

Age (Ma)

Pole

position

Turkmenbasi

reference

ref. on

Fig.6

lat.

long.

paleolat. dpal.

n°

Alborz (1)

36.5°N-5I.5°E

250-270

23°N

282°E

1 1

<B)

Abadeh and Jamal (1)

33°N-57°E

250-270

35°N

307°E

1 i

1 1

(A)

Mangyshlak (K)

44°N-52°E

240-250

61 °N

208°E

<D)

Kizil-Kaya (T)

4I°N-55°E

245-250

46°N

160°E

M 198 1

Turkeinia (T)

41°N-55°E

241-245

45°N

128°E

11.9

16.2

M 1981

Kizil-Kaya (T)

40.5°N-55.5°E

240-270

34°N

166°E

(E)

Sorkh Shales (1)

33.3°N-57.3°E

241-245

22°N

326°E

(Bj

Turkmenbasi (T)

40°N-53°E

193-227

15 4

(E)

Alborz (1)

35.7°N-52.3°E

178-235

49°N

I53°E

(B)

Shemshak (1)

33°N-57°E

185-200

38°N

3I4°E

(A)

Naiband and Nakhlak

33°N-57°E

160-200

I5°N

349° E

(A)

Chaloi Beds (T)

40°N-54°E

157-178

76°N

225°E

2.4

3,8

M 1016

Turkmenia (T)

4I°N-56°E

157-164

74°N

I09°E

2.7

3.3

47.3

M 1016

Garedu (1)

34°N"57°E

145-155

79°N

217°E

1 1

<B)

Kopel Dagh (T)

"Middle"

Cretaceous

75°N

I53°E

5.4

35.6

(F)

Alborz (1)

36.3°N-5L8°E

Cretaceous

61°N

147.5°E

6.1

9.3

32.1

(C)

Turkmenbasi (T)

40°N-53°E

37.6

(E)

Source

MARIE M. LEMAIRE ET AL.

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News, 1: 3-4.

Zijderveld, J.D.A., 1967.— Demagnetization: analysis of results. In: D. W. Collinson, K.M. Creer & S.K. Runcorn (eds),

Methods in Paleomagnetism: Proceedings of the NATO Advanced Study Institute on Paleomagnetic Methods, held in

the Physics Department of the University of Newcastle upon Tyne, April I-10, 1964. Elsevier, Amsterdam: 254-286.

Zonenshain, L.P., Kuzmin, M.I. & Natapov, L.M., 1990.— Geology of the USSR: a plate-tectonic synthesis. Geodynamic

Series , 21: 1-242.

Reconstruction of paleostress fields in Crimea and the

North West Caucasus, relationship with major

structures

Aline SAINTOT, Jacques ANGELIER, Alexander Ilyin

& Oleg Goushtchenko

Laboratoire Tectonique Quantitative, URA CNRS 1759, T26-E1, Case 129

University P. et M. Curie, 75252 Paris Cedex 05, France

ABSTRACT

Between the East European platform and the Black Sea domain, major deformations occurred during the Mesozoic and

Cenozoic times along the Crimea-Caucasus zone. Along the deformed boundary, especially in Crimea and NW-Caucasus, seven

tectonic events were identified based on inversion of fault slip data setsr The observed chronological criteria allowed us to

reconstruct the following succession of paleostress fields from youngest to oldest: (VII) Seismotectonic and tectonic studies

indicate a compressional stress regime for the Present, with o\ trending N-S or NNW-SSE in NW-Caucasus, NW-SE in

Crimea; (VI) a Early Pliocene NW-SE to WNW-ESE transpressional regime, with a counter-clockwise rotation of o\ trajectory

from NW-Caucasus to Crimea; (V) a nearly E-W transpressional regime, with a counter-clockwise rotation of o\ trajectory

from WNW-ESE in NW-Caucasus to ENE-WSW in Crimea, such as for the recent events; (IV) a Oligocene-Early Miocene

multidirectional extensional regime perpendicular to the axes of maximum subsidence in peripheric troughs of the chains; we

infer that these troughs developed during a post-compressional period following a lithospheric flexure. (Ill) a Late Eocene NE-

SW compression which produced the fold-and-thrust belt of NW-Caucasus; (II) a Paleocene event, with a transtensional regime

in NW-Caucasus related to the opening of the east Black Sea basin with 03 trending NE-SW. and a transpressional regime in

Crimea with o\ trending NW-SE; (I) the oldest event that we could identify, at the limit between Early and Late Cretaceous,

involves a N-S compressional regime in Crimea which resulted in north-vergent thrusting. We noted a deviation of stress

trajectories from NW-Caucasian chain to Crimean chain after the major compressional event (post Eocene). The development

of the N-S Kerch-Taman trough in the Oligo-Miocene time reflects the presence of a zone of weakness in the lithosphere, and

probably explains the fan-shaped trajectory patterns of the subsequent stress fields from NW-Caucasus to Crimea.

RESUME

Reconstruction des champs de contraintes en Crimee et dans le nord-ouest du Caucase, relation avec les structures

majeures.

Entre la plate-forme est-europeenne et le domaine de la Mer Noire, des deformations majeures ont eu lieu dans la zone

Crimee-Caucase au cours du Mesozoi'que et du Cenozoi’que. Le long de la frontiere deformee, plus specifiquement, de la

Crimee au Caucase nord-occidental, sept evenements tectoniques ont pu etre identifies par 1’inversion des mesures de failles a

stries. Les criteres chronologiques observes nous ont permis de reconstruire la succession suivante des champs de

paleocontraintes, du plusjeune au plus ancien : (VII) un regime actuel compressif determine par des etudes tectoniques et

Saintot, A., Angelier, J., Ilyin, A. & Goushtchenko, O., 1998.— Reconstruction of paleostress fields in Crimea and the

North West Caucasus, relationship with major structures. In: S.Crasquin-Soleau & E. Barrier (eds), Peri-Tethys Memoir 3:

stratigraphy and evolution of Peri-Tethyan platforms. Mem. Mus. natn. Hist, nat.. Ill : 89-112. Paris ISBN : 2-85653-512-7.

Source: MNHN. Paris

ALINE SAINTOT ET AL.

seismotectoniques, avec aj dirigt* N-S ou NNW-SSE dans le Caucase nord-occidental, et NW-SE en Crimee ; (VI) un regime

transpressif NW-SE a WNW-ESE au Pliocene inferieur. avec une deviation anti-horaire de la trajectoire de o\ du Caucase nord-

occidental a la Crimee ; (V) un regime transpressif, avec une deviation anti-horaire de la trajectoire de o\, comme dans le

regime actuel. de WNW-ESE dans le Caucase nord-occidental a ENE-WSW en Crimee : (IV) une extension multidirectionnelle

durant l’Oligocene et le Miocene inferieur ou l’axe 03 est perpendiculaire aux axes de subsidence maximale des depressions

peripheriques a la chaine. Ces fosses se developpent dans une periode post-compressive en reponse a la flexure de la

lithosphere. (Ill) une compression NE-SW a l'Eocene .superieur a I’origine des plis et des charriages des series de flysch du

Caucase nord-occidental ; (II) au Paleoc£ne, un regime transtensif dans le Caucase nord-occidental. relatif a I’ouverture du

bassin oriental de la Mer Noire ( 03 . oriente NE-SW), et un regime transpressif en Crimee (G\, oriente NW-SE) ; (I) une

compression N-S en Crimee. II s'agit de 1'evenement le plus ancien que nous avons pu identifie, a la limite Cretac 6 inferieur-

Cretace superieur, qui serait a l'origine des structures chevauchantes de la Crimee. Nous avons note une deviation dc la

trajectoire des contraintes du Caucase a la Crimee apres l' 6 venement majeur compressif de l'Eocene superieur. A I’Oligo-

Miocene, le d£veloppement du foss£ de Kerch-Taman entre les deux chaines selon un axe N-S serait relatif a la presence d’une

zone lithosph£rique fragilisee de meme orientation, qui perturberait les trajectoires de contraintes et expliquerait la distribution

en eventail des trajectoires de o\ sur la zone etudide.

INTRODUCTION

This study is located on the southern boundary of the east European platform, the Crimea and the

northwest Great Caucasus and surrounding areas. The general geology of the Crimea and the NW-

Caucasus has been studied since the end of the last century by works of numerous Russian and Soviet

authors. In rocks Late Jurassic to Neogene in age, we found evidences of brittle deformation. We

calculated local stress states by inversion of fault slip data sets. Groups of local stress states allow

reconstruction of paleostress fields which are closely related with development of major structures of the

deformed area. Moreover, the approach in term of paleostress reconstruction is an useful tool to

constrain the chronology of the deformation in the belt. In this paper, we aim at better understanding the

tectonic evolution of this area through identification of paleostress fields as indicated by brittle tectonic

analysis.

GEOLOGICAL FRAMEWORK AND AGE OF DEFORMATION EVENTS

The Crimean and NW-Caucasian mountains are classically described as two meganticlinoria

structures bordering the Eastern Black Sea Basin to the north. They correspond to the southern deformed

boundary of the Scythian plate, as far as the Paleozoic history is concerned. The zone corresponding to

present-day Crimea and Caucasus became part of Eurasia during the Namurian-Westphalian. During the

Mesozoic, the convergence between the Tethys and the Eurasian craton resulted in subductions and

collisions in the Crimean and Caucasian zones; the collision was continuing during the Cenozoic. The

structural framework of this region is summarized in figure la.

NW-CAUCASUS

The NW-Caucasus is an area of widespread folding and thrusting affecting flysch formations Early

Cretaceous to Paleocene in age. The fold axes trend NW-SE and the major thrusts associated are SW-

vergent (Fig. lb). In detail, some NW-SE faults perpendicular to the major folds and thrusts offset the

northern thrust fronts, which is also the case offshore to the south. The Oligocene formations

unconformably overlie the deformed flysch, indicating that the main folding event affecting the thick

flysch series is late Eocene in age.

Crimea

The core of the Crimean chain (Fig. 1c) is constituted by high karstic plateaus of Jurassic massive

limestones, forming thrusts sheets overthrusting series which include a parautochtonous flysch of the

Late Triassic-Aalenian (Taurian group, extremely deformed, see POPADYUK et al ., 1991), as well as

autochtonous series with ages up to the Albian. The general structure is shown in the cross-sections of

figure 2. The cross-section in the Indol River Valley (Fig. 2c) let us to observe an imbrication of

Valanginian-Hauterivian sandstones between to thrust planes to the feet of Agarmysh Mount. The recent

structural mapping of GALKIN et al (1994), in the area of the Salgir River Valley, reveals that

Source: MNHN. Paris

PALEOSTRESS IN CRIMEA AND NW-CAUCASUS

Structural scheme

of the studied deformed belt

and surrounding areas

[\^ Alpine

orogenic belts

i i Scythian plate

I | place of the foredeep

EE3

Southern boundary

of the east European platform

Regional fold trends

Contours of rift-faults

|-*-| Major Post-Eocene Thrusts

S Boundaries of maximum

subsidence areas

Fracturing in NW-Caucasus

[oilgocene-quatemary deposits

□ FlyschokJ fokJ-and-thrust zone

of NW-caucasian belt N

f^^j Syndine and anticline axis i

hrH Basin contour

INDOLO-KUBAN BASIN

Fault and thrusts

ikluch

BLACK SEA

PiG. 1.— Structural framework: a, general map; b. fracturing and structure in NW-Caucasus; c, fracturing and structure in

Crimea (from Geological map of the Great Caucasus, 1/500 000. anonymous, 198? and from Geological map of the

Crimea and Caucasus, 1/1 500 000, anonymous, 198?).

FlG. 1 .— Schema structural: a, carte generate: b, fracturation el structures dans le Caucase nord occidental; c, fracturation

et structures en Critnee (d'apres les cartes geologiques du Grand Caucase, 1/500 000, et de la Crimee, 1/1 500 000).

Source:

ALINE SAINTOT £T AL.

allochtonous units (Late Jurassic limestones and conglomerates) thrusted over the Early Cretaceous

series of the autochtonous complex. Late Cretaceous formations seal the thrust front (POPADYUK et al ,

1991) on the northern flank of the belt. A major phase of structuration in Crimea has thus occurred

between the Early Cretaceous and the Late Cretaceous (Fig. 2). The thrust sheets are north-vergent, with

a total displacement larger than 20 km (POPADYUK et al. y 1991).

Ml Chatyr Dag

Ml. Bayrakly

Aptian

Massive limestones

Sandstones, gritstones

Limestones and clays

Shales

Thrust

Fig. 2.— (from POPADYUK et al., 1991). A: cross-section along the Salgir River Valley; B: cross-section along the Tonas River

Valley (with sites 60 and 61); C: cross-section along the Sukhoy Indol River Valley (with site 76). Structures in

background in grey. Allochtonous complex: J 3 O-Km, massive conglomerates of Oxfordian-Kimmeridgian; J 3 . massive

tithonian limestones (with imbrications of Valanginian-Hauterivian shales and sandstones along the Sukhoy Indol River

Valley). Parautochtonous complex: T 3 -J 1 , terrigenous flysch of Taurian group. Autochtonous complex: Valanginian-

Hauterivian to Albian formations; C 2 -P 2 * Late Cretaceous-Eocene; N 2 . Late Neogene.

FlG. 2. — (d'apres POPADYUK et al., 1991).A : coupe de la vallee de la riviere Salgir ; B : coupe de la vallee de la riviere Tonas

(avec les sites 60 et 61) ; C : coupe de la vallee de la riviere Sukhoy Indol (avec le site 76). Structures en arriere plan

en grise. Complexe allochtone : J 3 O-Km, conglomerats massifs de VOxfordien-Kimmeridgien; Jj, calcaires massifs

tithoniens (avec imbrications d'argiles et de gres du Valanginien-Hauterivien le long de la vallee de la riviere Sukhoy

Indol). Complexe para-autochtone : Tj-Jj, flysch terrigene du groupe taurien. Complexe autochtone . formations du

Valanginien-Hauterivien a I'Albien ; C 2 -P 2 ’ Cretace superieur-Eocene ; N 2 . Neogene superieur.

Source: MNHN, Paris

PALEOSTRESS IN CRIMEA AND NW-CAUCASUS

The flyschs of the Taurian series (from Late Triassic to Aalenian) were strongly deformed during the

Earlier-Middle Jurassic phase of compressional deformation (Bathonian, NIKISHIN et al., this volume).

Concerning the Tertiary structure, outcrops of autochtonous terranes are present along the northern

side of the belt, gently dipping to the north. The Oligocene formations unconformably overlie the

Paleocene, indicating a late Eocene tectonic event also recognized in the NW-Caucasus. Three

unconformities are observed in Neogene time: the Karaganian horizon (Langhian, middle Miocene), the

Sarmatian formations (middle Miocene) and the Akchagylian series (late Pliocene) unconformably

overlie the older formations (for local stratigraphy, i.e. Eastern ParaTethys, see Table 1 in annex). They

are correlated with the period of maximum subsidence of basins (CHEKUNOV et al ., 1976; NIKISHIN et

al ., this volume). In contrast, the southern flank of the Crimean belt is cut by large faults with the

downthrown side to the south (KOGAN et al ., 1978; MURATOV et al ., 1984).

A major fracturating with dominating NE-SW and

N-S trends (Fig. lc) affects all terranes (including the

Late Neogene in Crimea). These major faults clearly

postdate the main thrusting discussed before, and are

oblique relative to the preexisting, nearly E-W

trending thrust faults.

Northeastern Black sea basin (Fig. 1)

The southern edge of the deformed belt is the East

Black Sea Basin (note that the Black Sea includes two

major basin, eastern and western, separated by a horst,

the Andrusov Ridge: KOGAN et al ., 1978; TUGOLESOV

et al., 1985). Numerous authors studied the

development of the-Black Sea through calculations of

subsidence curves based on analyses of well and

seismic reflection data.

Different interpretations exist concerning the

timing of development of the Black Sea basins. All

authors, however, concur to consider that the two

basins did not develop simultaneously, the western

basin appearing earlier than the eastern one

(ISMAGILOV et al , 1986; CHEKUNOV, 1990; OKAY et

a! ., 1994; NIKISHIN et al., this volume). The

geodynamic framework of the opening of the Black

Sea basins suggests the opening of a back-arc basin

related to a volcanic belt (the Pontides), as a response

of the north-vergent subduction of southern oceanic

domains related to the Tethys (LETOUZEY, 1977;

ZONENSHAIN et al. , 1986; OKAY et al. , 1994).

The Maykopian series (Oligocene-early Miocene)

unconformably overlie the earlier formations in basin

flanks, indicating a pre-Oligocene age for the

subsidence related to the main opening of the eastern

basin (ISMAGILOV et al., 1986). Accordingly, most

authors consider a Late Cretaceous to early Paleocene

age for the development of the East Black sea Basin

(OKAY et al. , 1994; ROBINSON et al. , 1996)

Note that the northwestern segment of the Black

Sea basin, south of Odessa shelf, is implied in north-

vergent thrusting, west of Crimea, where a seismic

reflection profile showed thrusts sealed by Oligocene

formations (POPOVICH, 1989). The northern part of the

shelf underwent reactivation during later tectonic

episodes. Especially, the East Black Sea Basin was

Table I.— Correlation of Eastern Paratethys Stratigraphy

with Tethys Stratigraphy (Neogene and Quaternary).

Table I .— Stratigraphic de la Para-Tethys orientate

(Neogene et Quatemaire).

0.50 Ma

Bakunian

0.73 Ma

Apcheronian

1 67 Ma

l—

<1)

Piacenzian

Akchagylian

3.40 Ma

Zanclean

Kimmerian

5.2 Ma

i—

Messinian

Pontian

7.0 Ma

Tortonian

Meotian

9.3 Ma

Sarmatian

Serravallian

13.7 Ma

Konkian

—

Langhian

Karaganian

Tchokrakian

16.5 Ma

Tarchanian

17.0 Ma

Burdigalian

Kozachurian

Sakaraulian

Caucasean

ALINE SAINTOT ET AL.

Table 2 — Summary of paleostress results at all sites studied. n°, reference number of the site. Stress regimes: S. strike-slip

faulting; R, reverse faulting; N, normal faulting. N. number of fault slip data. Trends and plunges of stress axes in

decrees (with *, back rotated stress tensor). Methods as referred to ANGELIER (1991). <t> ratio of stress differences,

<j >=(02 -< 73 )/(oi - 03 ). average angle between observed slip and computed shear in degrees. RUP, estimator as referred to

Angelier’(199I).Q, quality estimator (3, good; 2, fair; 1, poor). ...

TABLE 2.— Solutions calculees pour les sites etudies. n°, numero de reference du site. Regime de contratntes : S. decroclutnt:

R. inverse : N. normal. N, nombre de failles pour le calcul du tenseur des contraintes. Directions et plongements des

axes de contraintes en degres (avec *, tenseur des contraintes bascule). INVD ou R4DT: Methodes d'inversion

fANGELIER. 1991). ®=IO->-CTi)/(Oi -Oj). angle moyen entre la strie reelle et la strie calculee en degre. RUP. estimateur

moyen de qualitede la methode INVD (ANGEUER,' 1991). Q. critere de qualite (3, bon ; 2. moyen ; I, has).

moving to the NNW and plunged below the Crimean and NW-Caucasian belts, resulting in widespread

folding of the Maykopian series (ISMAGIL.OV et al., 1986). Thus, a south-vergent thrust zone developed,

which may indicate an incipient subduction context.

FOREDEEP BASINS

North of the belt, the E-W Indolo-Kuban basin probably represents the Maykopian foredeep basin of

the NW-Caucasus (NlKISHIN el al.. this volume), following the major compressional event, late Eocene

in age. Likewise, the Alma basin, north west of the Crimean, probably represents the mid-Tertiary

foredeep. Although these two basins are deep (Fig. la), large subsidence occurred during the Oligocene-

Miocene times along the northern side of the whole belt.

Southern troughs

Along the southern side of the Crimea-NW-Caucasus belt, the Sorokin and Tuapse troughs probably

represent flexural basins related to south-vergent thrusting. Especially, the Tuapse trough developed

along the northern slope of the eastern Black Sea Basin, parallel to the NW-SE trend of the major fold-

and-thrust structures of the NW-Caucasus, during the Oligocene-Miocene, as a consequence ol the

major compressional event of the late Eocene; similarly, the development of the Sorokin trough is

Oligocene-Miocene in age (TUGOLESOV et al., 1985, ISMAG 1 LOV et al., 1986; NlKISHIN et al.. this

volume). In addition, a late folding phase affected the Oligo-Miocene sedimentary infills of these

troughs (MEYSNER et al., 1981; ISMAGILOV et al., 1986; TEREKHOV et al., 1989).

Note that despite this relative regional continuity of the southern troughs, a transverse area ot

subsidence developed between the two folded areas (Crimea and Caucasus) with a N-S orientation: the

Kerch-Taman trough. Asymmetric mud volcanoes of Maykopian-Plio-Quatemary sediments developed

in the Kerch and Taman peninsulas, and in the Sorokin and Tuapse troughs as well. In the Kerch-Taman

areas, these mud volcanoes are elongated along E-W trends with steep flanks to the north indicating that

they formed at ramp fronts which correspond to the activation of deeper thrusts (KAZANTSEV et al.,

1988). Drilling in the Kerch peninsula revealed north-vergent thrusting of late Eocene age at depth of

about 3.5 km (KAZANTSEV et al., 1988). The diapirism of clay material, which resulted in later

volcanism, is related to later reactivations of these thrusts, under compressional regimes, and is now

active.

NlKISHIN et al. (this volume) point out that the subsidence phases related to the development of these

troughs was not accompanied by major normal faulting, which supports the interpretation in terms of

Mid-Cenozoic flexural basins. A geochemical study (AL'BOV et al., 1974) in the Kerch Taman rocks

revealed high concentrations of elements, deep crustal or mantellic in origin, which suggests that deep-

seated faults are involved, which is consistent with the probable presence of a crustal zone of weakness

beneath the Kerch-Taman segment of the belt.

Source: MNHN. Paris

PALEOSTRESS IN CRIMEA AND NW-CAUCASUS

Local ities

Lat /Long.

Age of

n°

Stress

CT 1

°2

°3

Meth.

in degrees

formation

regime

trend

plunge

trend

plunge

trend plunge

RUP

180

312

078

INVD

0.1

•

003

269

094

INVD

0.1

44.77/33.99

Middle Jur.

144

253

044

INVD

0.1

194

103

002

INVD

0.5

Dykes with E-W trends

44.72/33.85

Paleogene

100

288

190

INVD

0.6

129

031

220

INVD

0.4

44.77/33.87

Paleogene

131

294

041

INVD

0.5

44.77/33.98

Sant/Camp.

177

045

287

INVD

0.6

Crushed zone with N115 trends

44.78/33.75

Sarmatian

Tension gashes

a* 3 =

080

Tension gashes

150

Baschisarai

318

059

222

INVD

0.1

184

337

078

INVD

0.9

127

222

313

INVD

0.3

44.75/33.85

Late Eoc.

109

286

016

INVD

0.4

Tension gashes

<T' 3 =

045

Tension gashes

°3 =

145

44.73/33.92

Tur./Coniac.

Stylolites

a',= 150

44.80/34.02

Middle Jur.

Tension gashes

a* 3 =

080

Pre-tilting

Tension gashes

a' 3 =

120

44.80/34.01

Middle Jur.

Stylolites and associated stylolitic planes

ct',= 170

190

100

006

INVD

0.6

171

345

080

INVD

0.4

44.80/34.00

Middle Jur.

112

224

018

INVD

0.5

135

225

016

INVD

0.4

340

131

234

INVD

0.7

214

305

115

INVD

0.2

44.96/33.62

Neogene

•Tension gashes

a 3 =

180

Simferopol’

317

217

127

INVD

0.5

44.95/34.17

Middle Jur.

074

339

164

INVD

0.4

-xi-

178

311

- 0*70

— m —

080

VOi -

INVD

0.6

ITT

~~W~

44.41/33.70

Late Jur.

175

247

/ /

l l\>

103

343

INVD

0.5

246

149

045

INVD

0.4

44.37/33.68

Late Jur.

142

322

052

INVD

0.6

066

209

334

INVD

0.3

Foros

304

178

039

R4DT

0.2

44.37/33.68

Late Jur.

281

013

183

INVD

0.6

023

139

292

INVD

0.4

143

233

049

INVD

0.6

44.48/33.88

Late Jur.

130

270

001

INVD

0.3

44.42/33.82

Late Jur.

142

306

038

INVD

0.5

336

245

066

R4DT

0.1

283

044

134

INVD

0.3

Yalta

44.40/34.07

Late Jur.

340

090

180

INVD

0.5

252

064

161

INVD

0.6

202

053

294

INVD

0.4

44.57/34.07

Late Trias.

Early Mid. Jur.

023

215

124

INVD

0.5

44.49/34.02

Late Jur.

235

342

144

INVD

0.3

Stylolites

tj'j= 040

ALINE SAINTOT ET AL,

Localities

Lat./Long.

Age of

n°

Stress

CT I

a 2

Mcth. <I>

in degrees

formation

regime

trenc

plunge

trend

plunge

trend

plunge

RUP

44.52/34.22

Late Trias. Early

112

005

203

INVD 0.6

Jut.

Yalta

Stylolites

o' 1 — 150

44.53/34.03

Late Jur.

Stylolites

o’,= 040

Tension gashes

o’ 3 =

090

44.52/33.98

Late Jur.

348

148

257

INVD 0.6

44.57/34.33

Middle Jur.

253

073

343

INVD 0.3

Partenit

44.60/34.33

Middle Jur.

Pre-tilting tension gashes

& 3 = 100

44.62/34.32

Middle Jur.

180

000

090

INVD 0.4

45.06/34.62

Midldle Eoc.

Tension gashes

o' 3 = 060

45.07/34.67

Midldle Eoc.

297

177

087

INVD 0.3

216

054

314

INVD 0.7

Belogorsk

44.85/34.67

Late Jur.

197

097

305

INVD 0.3

44 87/34.63

Neocomian

347

171

078

INVD 0.7

44.97/34.63

Tithonian

164

066

254

INVD 0.4

008

142

274

INVD 0.4

45.05/34.60

Early Mid. Eoc.

327

141

234

R4DT 0.8

282

076

182

INVD 0.3

44.97/34.98

Late Jur.

208

048

310

INVD 0.3

359

258

103

INVD 0.3

Stari Krim

309

195

103

INVD 0.5

338

212

075

INVD 0.4

45.00/35.07

Cretaceous

Associated stylolites

014

258

106

INVD 0.5

247

156

021

INVD 0.1

Associated stylolites

148

263

357

INVD 0.6

•

107

265

010

INVD 0.6

44.80/34.94

Late Jur.

182

273

015

INVD 0.2

Tension gashes and associated mineral fibers 0 * 3 * 120

Tension gashes and associated mineral fibers

03 = 45

44.79/34.92

Late Jur.

Tension gashes

°3 =

155

Tension gashes

o- 3 =

020

Sudak

44.87/35.08

Late Jur.

276

084

182

INVD 0.7

Tension gashes

o*3*

100

44.86/35.07

Late Jur.

218

128

351

INVD 0.6

Tension gashes

0*3=

105

44.85/35.08

Late Jur.

172

047

270

INVD 0.5

129

006

273

INVD 0.4

44.84/35.08

Late Jur.

330

152

062

INVD 0.7

44.86/35.08

Late Jur.

293

151

061

INVD 0.1

44.90/35.13

Late Jur.

201

028

297

INVD 0.6

Feodosya

44.93/35.18

Middle Jur.

213

303

039

INVD 0.5

44.95/35.27

Late Jur.

018

22 01

156

281

INVD 0.6

•

017

111

287

INVD 0.6

45.32/36.34

Neogene

124

033

290

INVD 0.5

160

251

004

INVD 0.5

047

138

312

INVD 0.6

Kerch

45.37/36.26 Late Mid. Mioc.

212

303

093

INVD 0.4

45.40/36.42

Meotian

278

106

009

INVD 0.4

354

084

182

INVD 0.6

45.32/35.67

MeotVPont.

028

181

278

INVD 0.7

Lenino

45.40/35.73

Meotian

214

001

091

INVD 0.3

45.27/36.02

Middle Mioc.

068

175

338

INVD 0.6

Source: MNHN, Paris

PALEOSTRESS IN CRIMEA AND NW-CAUCASUS

Localities

Lat./Long.

Age of

n°

Stress

°l

a 2

CT 3

Meth.

in degrees

formation

regime

trend

plunge

trend

plunge

trend

plunge

RUP

174

031

299

1NVD

0.6

Late

•

288

029

119

INVD

0.6

44.87/37.33

Maas.

Tension gashes 03=

005

Anapa

Tension gashes 03=

120

44.92/37.40

Dan i an

Tension gashes 03=

135

45.02/37.47

Early

285

187

025

INVD

0.7

Paleocene

•

122

289

031

INVD

0.9

44.70/37.86

Late

202

298

028

INVD

0.5

Cretaceous

•

214

122

348

INVD

0.5

Tension gashes 03=

165

Tension gashes 03=

070

44.67/37.56

Early

250

354

152

INVD

0.3

Paleocene

249

350

156

INVD

0.3

052

296

202

INVD

0.4

067

291

198

INVD

0.4

218

308

045

INVD

0.5

•

038

308

152

INVD

0.5

Novorossisk

44.72/37.60

Maas?

359

240

111

INVD

0.4

•

195

343

104

INVD

0.4

Late

212

309

104

INVD

0.5

44.75/37.75

Cretaceous

029

206

296

INVD

0.2

218

039

309

INVD

0.3

44.88/37.90

Paleogene?

Tension gashes 03=

120

334

164

073

INVD

0.6

033

153

244

INVD

0.4

44.82/37.93

Coniacian

Pretilting tension gashes

ct' 3 = 090

Santonian

112

283

022

INVD

0.4

Pretilting stylolites and associated stylolitic planes o'|=

045

034

•232

135

INVD

0.6

44.83/37.07

Coniacian

033

271

125

INVD

0.6

Santonian

158

248

051

INVD

0.3

257

148

058

INVD

0.4

355

263

131

INVD

0.4

•

176

268

057

INVD

0.5

338

181

069

INVD

0.7

44.53/38.15

Late

329

061

200

INVD

0.8

Campanian

137

228

344

INVD

0.7

Stylolites and associated stylolitic planes

033

301

158

INVD

0.7

214

305

044

INVD

0.6

252

147

057

INVD

0.5

Gelendjik

44.55/38.17

Coniacian

Stylolites and associated stylolitic planes o’|= 110

Santonian

202

295

036

INVD

0.8

206

116

023

INVD

0.7

Associated bedding plane fault

44.54/38.15

Coniacian

165

256

036

INVD

0.2

Santonian

Associated bedding plane fault

132

280

042

INVD

0.4

Associated bedding plane fault

44.55/38.13

Coniacian

280

186

027

INVD

0.7

Santonian

Associated sty lolites

44.57/38.10

Late

102

290

138

020

INVD

0.6

Campanian

102

Synfolding bedding plane fault during NE-SW compressional event

ALINE SAINTOT ET AL.

Localities

Arkhipo-

Osipovka

Lat./Long.

in degrees

Age of

formation

Stress N fT|

regime _ trend plunge

CT 2

trend plunge

°3

trend plunge

Meth. a

% Q

RUP

44.47/38.40

Coniacian

Santonian

194

307

099

INVD 0.6 15 32 1

288

44.46/38.38

Late

Cretaceous

R 17 089

019

115

INVD 0.1 11 30 3

182 11 305 71

Associated stylolites

INVD 0.2 10 28 3

44.52/38.32

Late

Campanian

30 108

217

018

INVD 0.2 17 37 3

44.38/38.53"

Coniacian

Santonian

13 140

050

231

INVD 0.5 10 22 2

024

118

283

INVD 0.1 10 29 1

10 272

006

097

INVD 0.4 06 12 2

44.25/38.85

Late

Maas.

120

11 143

287

018

INVD 0.4 06 27 3

44.27/38.82 Late Maas.

44.28/38.79 Late Camp.

120

72T

122

14 143

052

294

INVD 0.7 12 29 2

15 002

7 321

093

189

219

056

44.30/38.77

Late Cret.

123

16 159

344

249

124

44.33/38.70 Late Maas.

331

155

217

287

092

056

INVD

0.3

124

087

326

225

109

16 015

144

285

44.17/39.22

Late

Campanian

109

336

182

066

109

202

152

010

359

103

268

INVD

0.3

0.4

27 190

334

100

INVD 0.8 10 24 2

INVD 0.2 12 39 1

INVD 0.2 10 30 3

INVD 0.3 18 51 1

INVD 0.4 19 39 3

INVD 0.4 13 41 2

INVD 0.2 12 34 3

110

219

008

101

INVD

0.5

44.15/39.18

Coniacian

•

016

187

278

INVD

0.6

Santonian

110

338

215

071

INVD

0.3

•

158

048

249

INVD

0.2

Coniacian

111

031

124

225

INVD

0.7

44.14/39.14

Santonian

•

224

134

334

INVD

0.6

Pretilting stylolites and associated stylolitic planes a j-

040

44.12/39.05

Earlv Paleoc.

112

068

320

230

INVD

0.4

44.14/39.05

Late Camp

113

172

268

068

INVD

0.3

Early Maas.

347

228

077

INVD

0.3

114

081

210

344

INVD

0.5

114

229

136

045

INVD

0.4

44.15/39.02

Early

114

009

209

115

INVD

0.3

Paleocene

015

266

107

INVD

0.2

114

096

195

353

INVD

0.6

091

308

183

INVD

0.6

115

130

230

036

INVD

0.4

0 /

44.16/39.02

Late Maas.

115

084

178

335

INVD

0.8

0 /

163

340

071

INVD

0.8

116

158

358

252

INVD

0.7

116

108

237

007

INVD

0.5

•

285

021

194

INVD

0.4

44.17/39.00

Early

116

356

207

112

INVD

0.4

Paleocene

•

016

186

285

INVD

0.6

116

248

151

000

INVD

0.8

•

195

288

018

INVD

0.8

44.17/38.97

Earlv Paleoc.

117

081

220

346

INVD

0.2

44.18/38.92

Late

118

074

304

172

INVD

0.3

Cretaceous

•

287

017

195

INVD

0.3

44.19/38.89

Late Maas.

119

250

114

019

INVD

0.6

44.61/39.08

Early Paleoc.

103

199

335

108

INVD

0.4

Goryachikluch

44.55/38.98

Valanginian?

105

331

125

216

INVD

0.4

•

209

300

069

INVD

0.4

44.35/39.23

Valanginian?

106

Tension gashes and associated mineral fibers

o' 3 = 080

Shaumyan

44.30/39.30

Late Camp.?

107

278

130

037

INVD

0.8

44.28/39.28

Late Cretac.

108

058

297

201

INVD

0.4

Source: MNHN, Paris

PALEOSTRESS IN CRIMEA AND NW-CAUCASUS

PALEOSTRESS FIELDS

Methods for paleostress field determination

Numerous brittle structures have been observed. By inversion of faults slip data (ANGELIER, 1975,

1984, 1991), we determined both the attitude of the three principal stress axes (ai, 02 , 03 , with

oi^G 2 >G 3 and pressure counted positive) and the value of the ratio between principal stress

differences: <J> = ( 02 - 03 )/( 0 i- 03 ). In each site, we separated subsets of fault slip data consistent with

different stress tensors. The quality of each determination was determined based on the number of data,

on the angle between the calculated shear and observed striae for each fault, and on the estimators 'RUP'

(as defined by ANGELIER, 1991). Note that as the misfit increases, a ranges from 0° to 180° whereas

'RUP' ranges from 0° to 200°.

For the regional synthesis, local stress states were grouped based on consideration of both the type-

orientation of the corresponding regimes (extensional, compressional or strike-slip) and the

chronological constraints. We thus reconstructed seven successive paleostress fields which are

consistent in terms of regimes and timing (SAINTOT el at ., 1996).

In detail, relative chronology between successive paleostress were obtained by using criteria such as

successive striations on a fault plane, crosscutting relationships, reactivation of conjugate fault planes.

Consideration of the relative ages of faulting and folding was of large help, because it allowed

distinction of pre-folding, syn-folding and post-folding paleostresses. Compilation of such relative

chronology criteria brought constraints on regional successions of events through correlation of local

results. The stratigraphic ages of affected terranes also imposed major constraints in dating.

REGIONAL RESULTS: CRIMEA AND NW-CAUCASUS BELTS

The studied region was divided in three parts based on regional geology, with 69 sites in the Crimean

chain, 11 sites in Kerch peninsula and 44 sites in the NW-Caucasus. 140 local stress states were

determined, and their characteristics are listed in table 2. The regional synthesis allowed characterization

of seven main paleostress fields. We describe them thereafter, from the latest to the oldest. Note that in

some cases, the tectonic regimes in Crimea may differ from those in NW-Caucasus, as the result of large

differences in the orientation of structural grains.

THE PRESENT-DAY STRESS FIELD (Fig. 3a, 3b, 3c).— The southern boundary of the east-european

platform is seismically active, so that the present-day stress field could be determined through inversion

of focal mechanisms of earthquakes (GOUSHTCHENKO et al., 1993a, b). In the studied area, the direction

of the principal horizontal stress axes widely varies (Fig. 3a). From the westernmost part of Crimea to

the eastern part of the NW-Caucasus, the trend of 0 ( trajectories deviates from NW-SE to N-S, the

attitude of 02 axes is E-W trending in the NW-Caucasus and vertical in western Crimea, while 03 axes

are SW-NE in Crimea and vertical in the NW-Caucasus. We conclude that the Crimean belt is presently

dominated by a strike-slip regime, whereas the NW-Caucasus undergoes a pure compressional regime

(Fig. 3a).

Concerning geological data (Fig. 3b, 3c), in the Kerch peninsula, we collected slip data on reverse

faults (sites 12, 78; Fig. 3c) associated to a pure compressional regime like in NW-Caucasus (sites 85,

92, 96, 121; Fig. 3b). This Quaternary regime affects Crimea where strike-slip faults were observed

(sites 1, 3, 41, 48, 61, 63, 79, 84). We include in this tectonic phase stress states which show E-W

extension (sites 25, 28, 30, 71, 86 , 106). This may be due to a release of confining pressure for the

compressional N-S event, or to a permutation of o 1 and C2 axes under the strike-slip regime that we

observed in sites in Crimea (sites 1, 3, 41, 48, 61,63, 79. 84); we note that in Crimea this phenomenon is

localized on the northern flank of the belt, with development of numerous tension gashes (sites 23, 25,

28).

Note also that the reactivation of the thrusting defined of late Eocene in age and which produces the

100

ALINE SAINTOT ETAL.

recent mud volcanism in the Kerch-Taman peninsula could be related to this compressional phase

(ANGELIER et al., 1994; SAINTOT et al., 1995).

The early Pliocene (Kimmerian) compressional stress field (Fig. 4a, 4b).— As the

Quaternary event described above, we observed a deviation of the trends of a] trajectories from NW-SE

in NW-Caucasus to WNW-ESE in Crimea (Fig. 4a, 4b).

The strike-slip regime dominated in NW-Caucasus (sites 14, 85, 92, 115, 116, 122, 123) but reverse

faulting was observed locally (sites 89, 120). This event affected, with formation of reverse and strike-

slip fault sets, the Neogene series in the Kerch area in sites 12,78 (ANGELIER et al., 1994). In the

Crimean belt, the strike-slip regime was observed in six sites (2, 5, 10, 41, 52, 76); one site (47) was

marked by pure compression southwest of Crimea. Development of tension gashes was associated with

this event (sites 6, 17) and near Bashisarai (site 25) extension occurred perpendicular to the trend of

compression.

THE nearly E-W compressional stress FIELD (Fig. 5a, 5b).— Along the NW-Caucasian coast

from Gelendjik to Tuapse, we observed development of reverse faults (sites 93, 88) and strike-slip faults

(sites 87, 89, 95, 102, 114, 117) in the Cretaceous to Paleocene series relative to a nearly E-W

compressional event (Fig. 5a). In this tectonic setting, the NW-Caucasian margin underwent a sinistral

strike-slip movement.

In the western Crimean belt, in Jurassic limestones, a strike-slip regime with c\ trending WSW-ENE

occurred (sites 8, 9, 32, 43, 45, 51), with local development of stylolites. In eastern Crimean belt, this

compressional regime is poorly recorded (site 79). In the Neogene series of the Kerch Peninsula (sites

12, 70, 74, 75), we observed a local deviation in the trend of o \ trajectories, from WSW-ENE to NE-

SW. In this tectonic setting, the Crimean boundary probably underwent a sinistral strike-slip movement.

The age of this event is attributed to the Sarmatian; unconformities observed in Crimea and NW-

Caucasus'and the period of subsidence defined by NIKISHIN et al. (this volume) support the age of this

event.

THE MULTIDIRECTIONAL EXTENSIONAL STRESS FIELD (Fig. 6 a, 6 b).— This extensional regime

probably occurred from Oligocene to early Miocene time (Maykopian), in agreement with the

development of peripheric basins of the Crimean and NW-Caucasian belt. The trend of a 3 trajectory

was thus perpendicular of the axes of maximum subsidence in the troughs.

Along the NW-Caucasian coast (Fig. 6a), Cretaceous and Paleocene series are affected by NE-SW

extensional regime related to the development of the Tuapse trough (sites 85, 96, 112, 114, 119, 120,

124). A NE-SW extensional regime is also observed in the northern flank of the NW-Caucasus (sites

107, 108), in agreement with the development of the Kuban foredeep (Fig. 6a).

A NW-SE extensional regime related to the development of the Alma trough is marked in the

Fig. 3.— Present-day Quaternary stress fields; a, computed by inversion of focal mechanisms of earthquakes (from

Gol'SHTCHENKO et at., 1993a, b); b and c. computed by inversion of fault slip data sets in NW-Caucasus and Crimea

respectively. Stereoplots in figure 3b-c: examples of faults consistent with the paleostress field illustrated. Schmidt’s

projection, lower hemisphere. Fault planes as thin lines, bedding planes as doted lines, striae as small arrows (inward

directed=reverse, outward directed=normal. strike-slip striae as couple of thin arrows). Computed stress axes as 5-, 4-

and 3- branch stars (g\, 02 and 03 respectively). Trend of compression: inward directed large arrows; trend of

extension: outward directed Targe arrows. Couples of arrows in map: sites where we reconstructed the tectonic regime

illustrated. Trend of compression: inward directed arrows; trend of extension: outward directed arrows; trend of G\ axis

in strike-slip regime: inward directed triangles. Brittle features associated with faults: JX, tension gashes; LP, stylolites;

JP. stylolitic pressure-solution seams. (For locations, see figures lb and Ic for NW-Caucasus and Crimea respectively).

FIG. 3 .— Champ de contraintes quaternaire et actuel ; a, determine par I'inversion des mecanismes an foyer des seismes

(d'apres GOUSHTCHENKO et at., 1993a. b); b et c. determine par I'inversion des glissements sur les plans de failles

respectivement dans le Caucase nord-occidental et en Crimee. Stereogrammes (Fig. 3b-c): exemples de failles

compatibles avec le champ de contraintes illustre. projection de Schmidt hemisphere inferieur. Plans de failles : lignes

continues, plans startigraphiques : lignes en pointille. stries : petites fleches (convergentes = inverses, divergentes =

normales, stries decrochantes: couples de fleches). Axes de contraintes calcules : etoiles a 5-. 4- et 3- branches. Oj. 02

et (7? respectivement. Direction de compression : grandes fleches convergentes; direction d 'extension : grandes fleches

divergent£s. Couples de fleches sur les cartes : sites ou le regime tectonique a etc determine. Direction de compression

= fleches convergentes ; direction d'extension- fleches divergentes ; direction de I'axe Gj dans les regimes

decrochants = triangles convergents. Structures cassantes associees aux failles : JX, fentes de tension : LP. stylolites ;

JP. plans stylolitiques. (pour la localisation, voir figure lb pour le Caucase nord occidental et le pour la Crimee).

Source: MNHN. Paris

FALEOSTRESS IN CRIMEA AND NW-CAUCASUS

a 1 trajectory

°2 trajectory

. 03 trajectory

I | Foredeep

1 1 Alpine orogen

5 East European platform

102

ALINE SAINTOT ET AL.

Fig. 4.— Kimmerian (early Pliocene) stress field; a, in NW-Caucasus; b, in Crimea; symbols as for figures 3b-c.

Fig. 4.— Champ de contraintes cimmerien (Pliocene inferieur); a, dans le Caucase nord occidental; b, en Crimee ; memes

symboles que pour les figures 3b-c.

Source: MNHN . Paris

PALEOSTRESS IN CRIMEA AND NW-CAUCASUS

103

BLACK SEA

Fig. 5.— Sarmatian (middle-late Miocene) stress field; a, in NW-Caucasus; b, in Crimea:

Fig. 5.— Champ de contraintes sarmatien (Miocene moyen-superieur); a, dans le Caucase nord occidental; b, en Crimee ;

mernes symboles que pour les figures 3b-c.

Source: MNHN. Paris

104

ALINE SAINTOT ET AL.

Fig. 6 .— Oligocene-early Miocene stress field; a, in NW-Caucasus; b, in Crimea; symbols as for figures 3b-c.

Fig. 6 .— Champ de contraintes de I'Oligocene- Miocene inferieur; a, dans le Caucase nord occidental; b, en Crimee ; memes

symboles que pour les figures 3b-c.

Source: MNHN. Paris

PALEOSTRESS IN CRIMEA AND NW-CAUCASUS

105

Fig. 7.— Late Eocene stress Field; a, in NW-Caucasus; b, in Crimea; symbols as for figures 3b-c.

FlC. 7.— Champ de contraintes de l'Eocene terminal; a, dans le Caucase nord occidental; b, en Crimee ; memes symboles

qae pour les figures 3b-c.

Source

106

ALINE SAINTOT ET AL.

northwestern flank of the Crimean belt (sites 25. 32, 39), with sets of normal faults and tension gashes.

The trajectories determined along the south coast of the Crimea follows the curve ol the Sorokin

trough from N-S near Foros to NW-SE from Yalta to Sudak (Fig. 6b). The subsidence of the southern

flank of the Crimean mountain, along large step-like normal faults (illustrated in figure 10), is

contemporaneous with this tectonic framework. ...

The origin of the development of the peripheric trough may lie in an important flexuring process

affecting the lithosphere (NiKISHIN et al., this volume) after the late Eocene compressional event

described below. .

We found some traces of an E-W extensional event (not illustrated) relative to the formation of the

N-S Kerch-Taman trough (SAINTOT et al., 1995) but this trough is a branch between the Indolo-Kuban

trough to the North and the Sorokin trough to the South of Crimea (Fig. la). The existence of a N-S

oriented weak lithospheric zone is inferred, and might have permitted propagation of the Sorokin trough

to the North and the propagation of the Indolo-Kuban trough to the South.

THE NE-SW COMPRESSIONAL STRESS FIELD (Fig. 7a, 7b).—In NW-Caucasus, the NE-SW

compression (ANGEL1ER et al., 1994) is very well marked in the series of flysch from Cretaceous to

Paleocene (sites of Paleocene formations: 14, 103). The characteristics of this compressional event (Fig.

7a) are in agreement with the general direction of thrusting and folding (SAINTOT et al., 1995). This

event occurred:

— before folding in sites 13, 14, 105, 111. with development of reverse faults sets, and in sites 15,

96, 114, 116 with strike-slip faults;

— syn-folding, with associated bedding slips (site 92);

— after folding in sites 16. 91 with reverse fault system, and in sites 14, 86, 87, 103, 109, 110 with

associated conjugate strike-slip faults.

Near Gelendjik (sites 95) and near Tuapse (sites 111), this event resulted in the development of

stylolites and associated stylolitic pressure solution seams before tilting.

The northwestern tip of the NW-Caucasian belt was affected, during this tectonic event, by an

extensional regime with <73 trajectories trending NW-SE (sites 14, 16, 97, 98). This local area may have

been affected by a release of the confining pressure within the frame of a general compressional regime.

The tectonic activity related to this major compressional event in NW-Caucasus affected Crimea with

a strike-slip regime (sites 30, 45, 47, 58, 60, 76, 79, 81). However, a permutation of 03 and <72 axes

occurred in site 79 , which produced a local extensional regime perpendicular in trend to the average

compression. Reverse faults (sites 5, 11,41. 68 ), tension gashes (site 79) and stylolites (site 58) locally

developed (Fig. 7b).

THE EXTENSIONAL-STRIKE-SLIP STRESS FIELD IN NW-CAUCASUS AND THE COMPRES-SIONAL-

STRIKE-SLIP STRESS FIELD IN CRIMEA (Fig. 8a, 8b, 8c).— An early NE-SW extensional-strike-slip event

affected flysch formation in NW-Caucasus at least up to the Early Paleocene epoch (formations of site

14 affected); this event occurred before the late Eocene folding phase. Strike-slip faults developed in this

tectonic setting (sites 14, 99, 114, 116), with a permutation of Gi and G2 stress axes occurring in site

116 (Fig. 8a).

From the Cretaceous to the Paleocene time, a deep trough with flysch deposits developed in the NW-

Caucasus, along NW-SE trends (which were folded in late Eocene compressional event), nearly parallel

to the east Black Sea basin (NiKISHIN et al., this volume). The extension was probably related to a

transtensional overall mechanism. The whole region was affected by an important strike-slip extensional

regime, and the interpretation of this event suggests that the opening of the east Black Sea basin

occurred until the early Paleocene.

A nearly E-W extensional regime is clearly defined in sites 109, 110, 115 and it is associated to a

strike-slip regime in sites 124. 113, 110, 109 (Fig. 8b). This event occurred before the folding event of

Fig. 8 .— Paleocene stress field: a, b in NW-Caucasus; c, in Crimea; symbols as for figures 3b-c.

FlG. 8— Champ de contraintes paleocene ; a. b dans le Caucase nord occidental; c, en Crimee : merries symboles que pour les

figures 3b-c.

Source . MNHN. Paris

PALEOSTRESS IN CRIMEA AND NW-CAUCASUS

107

108

ALINE SAINTOT ET AL.

late Eocene and it affected NW-Caucasian Paleocene flysch, so that it is difficult to distinguish this

event from the NE-SW extensional strike-slip discussed before (Fig. 8 a).

Simultaneously. Crimea was affected by a compressional strike-slip regime with a) trending NW-SE

(Fig. 8 c). In this event, reverse faults (sites 41.45. 47. 59) and strike-slip faults (sites 1. 2, 3. 9. 10, 15,

38," 47) developed. Normal faults were only identified in site 66 . Note the homoaxiality of stress in

Crimea and NW-Caucasus, and the change in strike-slip regime which was transpressional in Crimea

and transtensional in NW-Caucasus.

THE OLDEST N-S COMPRESSIONAL STRESS FIELD (Fig. 9).— In Crimea, the massive jurassic

limestones formed ihe high plateaus (Karabi Yayla, Chatyr Dag plateaus and others, Fig. 2) and are

strongly fractured along N-S directions. The N-S compressional event is recorded in these formations by

reverse conjugate fault systems (sites 1, 5. 40, 41). The compressional N-S event was associated with the

development of thrust sheet structures in Crimea. A strike-slip regime is associated to this event

(observed in sites 1, 32, 41,53, 63, 64, 76) with development of stylolites in site 76.

Near Bachisarai (sites 1. 41), in the northern-western flank of the Crimean chain, dykes of dioritic

rocks crop out, they are elongated in E-W direction, and shows a dip of 70° to the north. They formed in

Middle Jurassic time within the Tauric series (volcano-terrigenous flysch of the Late Triassic-Aalenian)

during an earlier N-S extensional event that we could not identify. We assume that these dykes

developed along vertical planes, indicating that they were later tilted during the N-S compressional

event (which produced strike-slip faults before tilting and reverse faults after tilting, with again a

permutation between G2 and <33 principal stress axes).

Concerning the age of the N-S compressional event, the affected limestones are Middle and Late

Jurassic in age. It is possible to assume that this event is related to the development of north-vergent

thrust sheet structures, in Crimea which occured after Early Cretaceous time, according to POPADYUK el

al. (1991) and GALKIN el al. (1994). Moreover, the Late Cretaceous limestones and clays seals the

thrust-front in the northern Bank of Crimea (POPADYUK et al., 1991). This tectonic event occured in the

limit between Early Cretaceous and Late Cretaceous. To confirm this point of view, it is necessary to

collect more fault slip data sets in Cretaceous formations in Crimea.

PlG. 9.— Early Cretaceous/Late Cretaceous stress field in Crimea; symbols as for figures 3b-c.

Fig. 9 .— Champ de contraintes Cretace inferieur/ Cretace superieur en Crimee ; mernes symboles que pour les figures 3b-c.

Source . MNHN. Paris

PALEOSTRESS IN CRIMEA AND NW-CAUCASUS

109

DISCUSSION AND CONCLUSION

Relations to major events

Summarizing, we reconstructed several paleostress fields related to major steps in the evolution of

the Crimea and the NW-Caucasus region. We present them below, in the stratigraphic order (Fig. 11):

— The N-S compressional event induced strike-slip faulting and north-vergent major thrusting,

resulting in structuration of Crimea in the Early Cretaceous/Late Cretaceous time (Fig. 9). No

paleostress state could be determined in the NW-Caucasus for this event.

— In the Paleocene time, a transtensional regime was active in NW-Caucasus region, with 03

trending nearly NE-SW. This regime was probably related to the rifting of the east Black Sea basin,

which is Late Cretaceous/Paleocene in age (Fig. 8 a, 8 b). In the Paleocene time, a transpressional regime

was active in Crimea with G\ trending NW-SE. This regime reactivated N-S directed fracturing in a

strike-slip regime in Crimea (Fig. 8 c).

— The late Eocene NE-SW compressional event resulted in major structural development of the

NW-Caucasus (Fig. 7a), with folding of the flysch mass and south-vergent thrusting of numerous units.

A transpressional regime with a\ trending NE-SW occurred in Crimea, and this margin underwent and

overall in sinistral strike-slip regime (Fig. 7b).

Fig. 10.— Photography illustrating the Oligocene-early Miocene event with a normal step-like fault near Foros (site 8) in the

southern coast of Crimea.

Fig. 10 .— Faille normale an miroir decametrique pres de Foros (site 8) sur la cote Sud de la Crimee i I lust rant Vevenement

oligocene-miocene inferieur.

ALINE SAINTOT FTAL.

1 10

Present-day

Quaternary

Crimea

100 km

-Caucasus

Kimmerian

(Lower Pliocene)

100 km

fsJW-Caucasus

\\\w

Reactivated fracturing

Folding and development of mud

volcanoes in the Kerch-Taman areas

Reactivated fracturing

Major thrusting in Crimea

Thrusts are E-W directed faults

N-S fracturing occured in Late

Jurassic limestones

Fig. 11 .— Tectonic evolution of the NW-Caucasus and Crimea in the light of successive reconstructed paleostress fields.

Directions of compression, extension and lateral shear shown as couples of arrows. For details, see figures 3, 4, 5, 6, 7,

8 , 9. , .... .

FlG. II .— Evolution tectonique du Caucase nord occidental et de la Crimee deduite de la succession des champs de

contraintes. Directions de compression, d'extension schematises par des couples de {leches. Pour plus de details, voir

les figures 3, 4, 5, 6, 7. 8, 9.

Source: MNHN. Paris

PALKOSTRESS IN CRIMEA AND NW-CAUCASUS

111

— A multidirectional extensional event related to the development of peripheric basins of the belt

occurred in the Oligocene-early Miocene times. The trend of a 3 trajectories was perpendicular to the

axes of maximum subsidence in the nearest troughs (Fig. 6 a, 6 b).

— A nearly E-W strike-slip regime, not very well constrained, seems to have been active in the

Sarmatian time (Fig. 5a, 5b). The NW-Caucasian margin underwent sinistral strike-slip movement; the

sinistral movement of the Crimean margin was unclear. A deviation in the trend of o\ trajectory

occurred, from E-W in the NW-Caucasus to ENE-WSW in the Crimea.

— The Kimmerian event is well defined and from Crimea to NW-Caucasus the margin was in a

general dextral strike-slip regime. The dextral transpressional reactivation of the belt probably uplifted

Crimea as a ‘push-up’ structure between the West Crimean Fault (Fig. la) and the NW-Caucasian

margin (SAINTOT et al. } 1995). We note that a deviation of the trend of o\ trajectories from NW-SE in

NW-Caucasus to WNW-ESE in Crimea (Fig. 4a, 4b).

— The Quaternary-Present Day stress field determined by inversion of seismic focal mechanisms

and recent fault slip data is characterized by a NW-SE tranpressional regime in Crimea, whereas a N-S

compressional regime prevailed in the NW-Caucasus (Fig. 3a, 3b, 3c). As for the previous Sarmatian

and Kimmerian events, we observed a counter-clockwise deviation in the trend of G] trajectories, from

the NW-Caucasus to the Crimea. This phenomenon is probably due to the presence of a N-S trending

zone of lithospheric weakness which separated the Crimea and the NW-Caucasus, as indicated by the

development of the N-S Kerch-Taman trough in Oligocene-early Miocene time.

The latest tectonic events (Sarmatian, Kimmerian and Quaternary-Present day) were marked by

unconformities both on flank of Crimean belt and NW-Caucasian mountains.

An important feature of the whole area studied is the large number of strike-slip faulting occurrences.

We infer that during the Alpine tectogenesis, the southern boundary of the southern part of the east

European Platform mainly underwent alternating transpressive and transtensive regimes.

AKNOWLEDGEMENTS

We thank Russian participants in the first field trip, and Ms Carla MULLER for the systematic dating

of rock samples. This work was supported by the Peri-Tethys program and thesis grants were supported

by the French Embassy in Moscow (A. ILYIN) and the French M.R.E. (A. SAINTOT).

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Galkin, V.A., Fedorov, Ye.V. & Bakhor. K., 1994.— The interrelationships and structure of the Upper Jurassic and Lower

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Oceanology, 26 (4): 486-491.

Kazantsev, Yu.V. & Bekher. N.I., 1988.— Allochtonous structures of the Kerch Peninsula. Geotectonics 22 (4): 346-355.

Kogan, L.I.. Malovtskiy, Ya. P.. Moskalenko, V.N. & Shimkus, K.M., 1978.— New data on structure of the sedimentary

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Nikishin A.M., Cloetingh, S.. Baraboshkin E.Yu.. Bolotov, S.N., Kopaevich, L.F., Nazarevich, B.P., IXnov, D.I.,

Brunet, M.F., Ershov. A.V., Kosova, S.S.. Il’Ina. V.V. & Stephenson, R.A., 1998.— Scythian platform:

chronostratigraphy and stages of tectonic histoi 7 . In: S. Crasquin-Soleau & E. Barrier (eds), Peri-Tethys Memoir 3:

stratigraphy and evolution of Peri-Tethyan platforms. Memoires du Museum national d'Histoire naturelle, 177 : 00-000.

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surrounding regions. Geology. 22 : 267-270.

Popadyuk, I.V. & Smirnov. S.Y., 1991.— The problem of the structure of the Crimean Mountain region: traditional ideas and

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Sea. Geotectonics, 23 (6): 530-537.

Robinson, A.G.. Rlidat. J.H., Banks, C.J. & Wiles, R.L.F., 1996.— Petroleum geology of the Black Sea. Marine and

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SAINTOT, A., ANGELIER, J., GOUSTCHENKO, O.. ILYIN, A., REBETSKY, Y.. VASSILIEV, N.. YAKOVLEV, F. & MALUTIN, S.,

1995.— Paleostress reconstruction and major structures in Crimea and NW-Caucasus. Communication to the EUG

VIII, Strasbourg, 9-13 April 1995. Abstract Supplement n°l to Terra Nova. 1: 269.

SAINTOT, A., ANGELIER J. & Ilyin, A., 1996.— Champs de contraintes en Crimee et dans le Caucase nord-occidentale: 16th

Earth Sciences meeting, 10-12 April 1996, Orleans. Societe geologique de France. Paris. Abstract in special volume.

Terekhov, A.A. & Shimkus K.M., 1989.— Young sediments and young overthrust structures in the Peri-Crimean and Peri-

Caucasus zones of the Black Sea Basin. Geotectonics, 23 (1): 54-59.

Tugolesov, D.A., Gorshkov. A.S., Meysner, L.B., Solov’Yev, V.V. & Khakhalev, Ye.M., 1985.— The tectonics of the

Black Sea trough. Geotectonics, 19 (6): 435-445.

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basins. Tectonophysics. 123: 181-211.

Source: MNHN. Paris

Neogene evolution of the Carpathian foothills:

insights from the romanian diapir fold area

Jean-Claude HlPPOLYTE 1 " & Mircea SANDULESCU ' 21

"' Laboratoire de Tectonique Quantitative, CNRS URA 1759. T26-E1. Case 129

University P. et M. Curie. 75252 Paris Cedex 05, France

,2 ' Institut Geologique de Roumanie, Str. Caransebes 1, Sector 1.78344 Bucarest, Roumanie

ABSTRACT

In the diapir fold area of Romania, the frontal part of the Carpathian foothills is characterized by folds and minor thrusts

formed during the Pliocene-Quaternary 'Wallachian phase’. Most Wallachian folds and thrusts are not located at the outermost

front of the Carpathians, but are more internal structures. Consequently, the location of the deformation during the 'Wallachian

phase' contrasts with the Cretaceous-middle Miocene piggy-back thrust sequence of the East Carpathians. This difference

results from a modification in the stress field, with a direction of compression changing during late Miocene from NW-SE to

NS. Structural observations and the stratigraphic dating of the paleostress change indicate that the Wallachian tectonics was not

a brief event but lasted from middle Tortonian to early Pleistocene, with major deformations at the end of this period. The

building of the inner margin of the foredeep was contemporaneous with the strong subsidence of this basin and with the filling

of few piggyback syn-tectonic basins.

RESUME

Evolution neogene du eontrefort des Carpates : aper^u de la zone des plis diapirs de Roumanie.

Dans la zone des plis diapirs de Roumanie, la partie frontale du eontrefort des Carpates est caracterisee par la presence de

structures compressives, formees durant la phase Wallache d'age Plio-Quatemaire. La plupart de ces structures ne sont pas

localisees sur le front externe des chevauchements miocenes. mais sont situees a l'interieur des nappes miocenes. La localisation

de la deformation pendant la phase Wallache contraste done avec la sequence d’apparition des chevauchements anterieurs au

Cretace-Miocene moyen des Carpates orientales. Cette difference est liee ici a une modification du champ de contraintes.

puisque la direction de compression est pass£e de NW-SE a N-S durant le Miocene superieur. La presence de plis

synsedimentaires et la datation stratigraphique des etats de contraintes successes montrent que la tectonique Wallache n’est pas

un evenement bref, mais a duree du Tortonien moyen au Pleistocene inferieur, avec une deformation majeure & la fin de cette

periode. La construction de la marge interne de 1'avanl-fosse actuelle fut contemporaine de la forte subsidence de ce bassin et du

remplissage de quelques bassins syn-compressifs de type "piggyback".

INTRODUCTION

The Carpathian fold and thrust belt results from the closure of the Tethyan oceanic basin and its

continental margins. The 'Main Tethyan Suture’ runs across the Carpathian area (Fig. 1). It separates

HlPPOLYTE, J.-C. & SANDULESCU, M., 1998.— Neogene evolution of the Carpathian foothills: insights from the romanian

diapir fold area. In: S. Crasquin-Sdleau & E. Barrier (eds), Peri-Tethys Memoir 3: stratigraphy and evolution of Peri-

Tethyan platforms. Mem. Mus. natn. Hist, nat ., 177 : 113-127. Paris ISBN : 2-85653-512-7.

Source: MNHN, Paris

JEAN-CLAUDE HIPPOLYTE & MIRCEA SANDULESCU

1 14

Fore-Apulian tectonic units (Inner Dacides) to the west from European units (Median, Outer and

Marginal Dacides plus Moldavides) to the east and southeast. The Dacides show Cretaceous

tectogenesis, while the Moldavides show Miocene tectonism (SANDULESCU. 1980; 1989). Since the age

of the formations of the thrust units and the age of their thrusting decrease from the internal to the

external areas (DUMITRESCU & SANDULESCU. 1970) there is a typical piggy-back sequence of thrusting

in this fold and thrust belt.

Each period of tectogenesis (Cretaceous and Miocene) was followed by flexural subsidence of the

foreland (ELLOUZ & ROCA, 1993). In front of the Cretaceous thrust belt, a migrating foredeep basin was

filled with Cretaceous to Miocene flyschs. These flysch sediments now constitute the moldavian nappes

emplaced during Miocene (SANDULESCU. 1980; 1989). The present foredeep basin which follows the

curvature of Carpathians (Fig. 1) is partly filled by middle Miocene sediments and by the outermost

Carpathian nappe. In the southern areas it also includes a thick late Miocene to Early Quaternary

sedimentary sequence, and it ranges 9000 metres depth in the Focsani Depression (DUMITRESCU &

SANDULESCU, 1970) (Fig. 1).

The outer part of this foredeep basin lies above the foreland Mesozoic carbonate platforms (Fig. 1)

and is largely covered by Late Quaternary eolian and fluviatile deposits. Its inner part is complex

FlG. 1.— Geological sketch of Romania, with location of the area of interest. 1. East European and Scythian platforms; 2, North

Dobrogea Cimmerian orogen; 3, Moesian platform; 4 . middle Miocene-Quaternary depressions; 5, Moldavides; 6,

Neogene volcanic arc; 7. Eocene-early Miocene post-tectonic cover;8. Median, Outer and Marginal Dacides (European

units); 9, Transylvanides and Pienides (Tethyan Suture); 10. Inner Dacides (Fore-Apulian units).

FlG. I.— Schema geologique de la Roumanie avec situation de la zone d'etude. 1. Plate-forme est europeenne et plate-forme

scythique ; 2. Orogene cimmerien de la Dobrogea septentrionale ; 3, Plate-forme Moesienne ; 4, Bassins du Miocene

moyen-Quaternaire : 5. Moldavides ; 6. Arc volcanique neogene ; 7. Couverture post-tectonique eocene-miocene

inferieur ; 8. Dacides medianes, externes et marginales (unites europeennes); 9, Transylvanides et Pienides (suture

tethysienne) ; 10. Dacides internes (unites pre-Apuliennes).

Source: MNHN. Paris

NEOGENE EVOLUTION OF THE CARPATHIAN FOOTHILLS

1 15

A outer front of shortening

l_ thrust

l— syncline

►- anticline

t- fold covered by sediments

r major Wallachian structures

Cretaceous thrust sheets

(outer Dacides)

Miocene thrust sheets

(Moldavides)

Meotian-Quatemary

sediments

direction of

compression

Buzau

Ploiesti

Valea Lunga Basin

foredeep

Fig. 2.— Directions of compression of Wallachian age (Meolian-early Pleistocene). Structural sketch from Dumitrescu &

Sandulescu (1970). The compression directions are dated oLthis period because they are computed with faults

measured in the Meotian-early Pleistocene sequence (sites 7, 10, 11. 12. 13, 14, 15, 16), or correlated using relative

chronologies (sites 8, 9 and 17), or correlated to an adjacent site with same trend of compression (sites 2 and 3).

FIG. 2.— Directions de compression d’dge Valaque (Meotien-Pleistocene inferieur). Schema structural d’apres DUMITRESCU <&

Sandulescu (1970). Les directions de compression sont datees de cette periode car elles ont ete calculees a partir de

plans stries mesures dans les depdts du Meotien-Pleistocene inferieur (sites 7. 10. 11. 12. 13. 14. 15. 16). ou pane

qu’elles sont correlees grace a des obsen'ations de chronologies relatives (sites 8. 9. 17), ou bien parce qu’elles sont

cor relees a des sites voisins mieux dates et qui presentent la meme direction de compression (sites 2 el 3).

because in places it is partly underthrust beneath the Carpathian thrust front (Fig. 1), whereas in other

places it partially rests above the compressional structures of the outer Carpathians (Figs 1, 2). In

northern Romania for instance, the Sarmatian deposits (Serravallian-early Tortonian of the eastern

Paratethys chronostratigraphy, figure 3, STEININGER et al„ 1988) of the foredeep margin are largely

overthrust by the frontal nappe of the Carpathians (Fig. 1). In the south, along the inner foredeep margin,

the late Sarmatian to Quaternary sediments of the foredeep sequence dip to the east, with sub-vertical

bedding planes close to the outcropping thrust units (Fig. 2). Still farther south, in the Carpathian bend

area of Romania, the late Sarmatian to Pleistocene foredeep sediments rest unconformably above the

outer Carpathian nappes (Figs 1, 2, 4). There, folds and minor faults of Wallachian age (early

Pleistocene, DUMITRESCU & SANDULESCU, 1968) characterize the diapir fold area (Figs 2, 4). Westward

(Getic Depression, Fig. 1), the foredeep of the South Carpathians has also largely covered the Miocene

nappes, but no Wallachian fold is described there, the Pliocene-early Pleistocene sediments ot the

foredeep gently dipping to the south.

To determine the way this complex margin was created, we made structural analyses supported by

fault-slip measurements primarily in the east Carpathians bend area (Fig. 2), where the Wallachian phase

was first defined (STILLE, 1924). Using the stress inversion technique (ANGE1.1ER, 1989), we will

tentatively relate the observed structures to their causative tectonic forces.

I 16

JEAN-CLAUDE HIPPOLYTE & MIRCEA SANDULESCU

Myr Eastern Paratethvs

Mediterranean

STRUCTURES

5.6

8.5

13.6

16.5

Pleistocene

Romanian

Pliocene

Dacian

Messinian

Pontian

Meotian

Tortonian

Sarmatian

Serravallian

Badenian

Langhian

In the East Carpathians bend area (Fig. 2),

seismic profiles show that the outermost thrust

sheet (Subcarpathian Nappe) is thrust above

Badenian and in places early Sarmatian deposits

(Figs 4, 5) (STEFANESCU, 1984; MOCANU &

RADULESCU, 1994; DlCEA, 1995). Being sealed

by late Sarmatian deposits (Fig. 5) (DUMITRESCU

& SANDULESCU, 1968), this thrusting ended in

Sarmatian. However, a Sarmatian to Quaternary

foredeep sequence rests above the outer

Carpathian thrust sheets (Fig. 5), and folds and

thrusts affecting sediments as young as early

Pleistocene exist also in this area (Fig. 4).

Therefore we can distinguish the pre-late

Sarmatian structures from the more recent ones

(Wallachian structures).

We can note in the cross-section of figure 4

that the amplitude of the Wallachian folds

increases from the outer to the inner foredeep,

where thrust faults even merge to the surface. It is

a general feature in the bend area, the most

important Wallachian folds and thrusts constitute

the compressive inner margin of the foredeep

basin (Fig. 2).

In figure 2, one can note that the Wallachian

structures (structures that affect the Meotian to

Quaternary sequence) are all located behind the

front of the Subcarpathian nappe. Taking into

account the age of this frontal thrust (intra-

Sarmatian), we conclude that the Wallachian

(Pliocene-Quaternary) thrusts are in overstep

sequence. In more details, seismic profiles show that in the late Sarmatian to Quaternary deposits, the

more internal the faults are, the younger they are (STEFANESCU & DlCEA, 1995).

The most important Wallachian thrusts, underlined on figure 2, constitute the compressive inner

margin of the foredeep. This margin extends from Valea Lunga to Buzau (Fig. 2) and trends E-W. It is

oblique to the front of the Subcarpathian nappe which trends NE-SW in this area (Fig. 2). In the cross-

section cf figure 4, this margin (area of the Valea Lunga basin) is more than 30 km internal to this front,

while in the cross-section of figure 6, this margin is only a few hundred metres internal.

Along the Focsani depression (Fig. 1), as mentioned before, the inner limb of the foredeep basin is

tilted to the east (Fig. 2). This regional tilt can result either from of a back shear motion of the late

Miocene-Pliocene molasse decoupled at the top of the Subcarpathian nappe (ELLOUZ et al ., 1996) (the

front of this nappe forming a triangle zone in a cross section view), or from the activity of an out-of¬

sequence east-vergent thrust fault (the Casim-Bisoca fault) (Fig. 6) (STEFANESCU, 1984; ELLOUZ et al .,

1996). In this later solution, we can note that the Casim-Bisoca fault has the same characteristics as the

thrust faults of the Valea Lunga-Buzau compressive margin. This faults is effectively in overstep

sequence compared with the front of the Subcarpathian nappe unconformably overlain by the late

Sarmatian to Pliocene sequence, and it constitutes the inner margin of the foredeep basin (Fig. 2). It

constitutes therefore the prolongation to the north of the Valea Lunga-Buzau compressive margin which,

north of Buzau, has the same trend as the front of the Subcarpathian nappe.

The obliquity of the Valea Lunga-Buzau compressive margin and the overstep sequence of thrusting

are the main characteristics of the foreland basin that we tried to explain.

Fig. 3.— Correspondence between Eastern Paratethys and

Mediterranean ages (from STEININGER el al., 1988).

FlG. 3 .— Correlation entre les ages de la Paratethys orientate

et ceitx de la Mediterranee (d'apres STEININGER et al..

1988).

Source: MNHN. Pans

NEOGENE EVOLUTION OF THE CARPATHIAN FOOTHILLS

117

FlG. 4.— Cross-section I, in the Wallachian fold and thrust area (location in Fig. 2; from STEFANESCU, 1984) with Schmidt

diagrams (lower hemisphere) of Wallachian striated fault planes (thin lines), computed stress axes (5-. 4- and 3- branch

stars represent Ol, 02 and 03 respectively) and bedding planes (broken lines). Site 14 is in the thrust contact between

Badenian (Langhian-early Serravallian) and Meotian sediments, h Meotian (middle Tortonian) to Quaternary foredeep

sediments; 2. diapir of early Miocene salt; 3. Oligocene-Sarmatian (middle-late Miocene); 4. Cretaceous-Eocene; 5.

autochtonous Miocene; 6, autochtonous Paleozoic-Paleogene.

FlG. 4— Coupe I. dans la zone des chevauchements et pi is Valaques (voir figure 2 pour la localisation ; d'apres STEFANESCU

1984) avec diagrammes (projection Schmidt, hemisphere inferieur) de plans de failles stries d'dge Valaque . axes de

contraintes calcules (etoiles a 5. 4 et 3 branches representant ol. o2 et 03 respectivement) el plans de stratification

(tirets). Le site 14 est silue au contact chevauchant entre les sediments d'dge badenien (Langhien-Serravallian

inferieur) et ceux d'dge meotien. 1, sediments d'avant-fosse d'dge meotien (Tortonien moyen) a Quaternaire ; 2. diapir

de sel miocene moyen ; 3. Oligocene-Sarmatien (Miocene moyen-superieur) ; 4. Cretace-Eocene ; 5, Miocene

autochtone ; 6, Paleozoique-Paleogene autochtone.

PALAEOSTRESS ORIENTATIONS

In an attempt to determine why the Valea Lunga-Buzau foredeep margin is oblique to the front of the

Carpathians and why the recent thrusts are in overstep sequence, we first measured fault planes with

slickensides and then computed the successive stress axes that account for the observed deformation.

Indeed two successive stress fields are reconstructed thus indicating NW-SE and N-S trends of

compression. The N-S trending compression (Fig. 2) is the most recent. It is the only compression found

in the Meotian to Quaternary foredeep sequence. Even the early Pleistocene sediments are affected (site

13, Figs 2, 4 and table 1). Therefore, this N-S trend of compression, identified in the main structures of

the inner foredeep margin, characterize the Wallachian phase (ST1LLE, 1924).

The stress regime is reverse (site 7, 10, 12, 13, 14 and 16, table 1) or strike-slip (sites 2. 3, 9. 11. 15

and 17, table 1) depending on the location (geographic stress permutation. HlPPOLYTE et ol., 1992).

Moreover, the existence of both reverse and strike-slip stress regimes in a same site (site 8 . table 1),

results from a temporal stress permutation (HlPPOLYTE el al ., 1992), probably during a unique event.

This N-S compression is sometimes perpendicular to the Pliocene-Quaternary folds axes, but it is

often slightly oblique, suggesting that some of them may be Moldavian structures reactivated during the

I 18

JEAN-CLAUDE HIPPOLYTE & MIRCEA SANDULESCU

CARPATHIAN FOREDEEP

SUBCARPATHIAN

NAPPE

MOESIAN PLATFORM

Fig. 5.— Seismic profile in the inner Carpathian foredeep (location in Fig. 2; from MOCANU & Radulescu, 1994, and Dicea,

1995). The frontal thrust of the Carpathians (Subcarpathian Nappe) is buried under foredeep sediments. Whereas the

tectonic contact between this nappe and the early foredeep sediments is sealed by Sarmatian deposits, the Pliocene

foredeep sediments are folded in the inner zones. PZ: Paleozoic, Tr: Triassic, J: Jurassic, K: Cretaceous, Ol: Oligocene,

Mi: Miocene, Z: Miocene salt. Bn: Badenian, Sm: Sarmatian, PI: Pliocene.

FlG. 5. — Profit sismique dans I'avant-fosse interne des Carpathes (voir figure 2 pour la localisation ; d’apres MOCANU &

RADULESCU, 1994 ; Dicea, 1995). Le chevauchement frontal des Carpathes (Nappe Subcarpatique) est enfoui sous les

sediments de Favant-fosse. Bien que les sediments Sarmatiens scellent le contact de cette nappe avec les depots

inferieurs de Favant-fosse, les sediments pliocenes de cette avant-fosse sont deformes dans les zones internes. PZ :

Paleozoique, Tr: Trias. J: Jurassique, K : Cretace, Ol: Oligocene, Mi: Miocene, Z: Sel miocene. Bn : Badenien, Sm :

Sarmatien, PI: Pliocene.

Wallachian phase (Fig. 2). This N-S compression also created the thrusts of Wallachian age bordering

the foredeep basin to the north (Fig. 2). Northwest of Ploiesti, the Valea Lunga depocentre is partly

isolated from the main foredeep basin (Fig. 2). This depocentre is located on the hanginwall of a thrust

fault. As measurements at site 14 (Fig. 4) indicate that this thrust moved to the south, the direction of

thrusting is found parallel to the direction of compression. The NE-trending segment of this thrust (Fig.

2) is an oblique ramp with a left lateral component of movement. The Valea Lunga basin, only deformed

by compression, is probably a piggyback syn-compressive basin (ORI & FklEND, 1984) dating both the

thrust activity and the N-S compression as Meotian-early Pleistocene in age. Thinning of the foredeep

sedimentary sequence over numerous anticline hinges in the East Carpathians bend area (Fig. 2),

confirms this age (STILLE, 1953; Paraschiv, 1975; STEFANESCU, 1984). In figure 6, the large thinning

of the Meotian-Quaternary sequence from the south to the north of a large broken anticline is a clear

example of syn-depositional tectonics.

In contrast to the foredeep sequence, the Paleogene to Sarmatian sediments of the outer Moldavian

nappes have recorded two compressional trends: N-S (Fig. 2) and NW-SE (Fig. 7). As the NW-SE

compression does not affect the Meotian to Quaternary sediments, we concluded that it is the older

Cenozoic event. Moreover, this stratigraphic dating of the stress orientation indicates that the NW-SE

compression ended in Sarmatian and was replaced at this time by the N-S compression, syn-depositional

of the Meotian-early Pleistocene stratigraphic sequence. Observations of slickensides superposition

Source: MNHN. Paris

NEOGENE EVOLUTION OF THE CARPATHIAN FOOTHILLS

119

Table 1.— Paleotress tensors computed from fault-slip data (Schmidt diagrams in annex). Sites of figs 2 and 7. S = State of

stress: S = strike-slip stress regime, C = compressional regime, E = extensional regime. N = number of faults used for

paleostress calculation. Ol, 02, 03: trend (N to E) and plunge of stress axes (°). O = (o2-o3)/(Gl-G3). ANG = average

angle between computed shear stress and observed slickenside lineation (°).

Tableau /.— Paleocontraintes calculus a partir de mesures de plans de failles stries (diagrammes en annexe). Sites des

figures 2 et 7. S= Type d’etat de contraintes : S = regime decrochant, C = regime compressif E = regime extensif N =

nombre de failles utilisees dans le calcul des paleocontraintes. ol, o2, o3 : direction et plongement en ° des axes de

contraintes. 0 = (o2-o3)/(ol-o3). ANG = angle moyen entre la contrainte cisaillante calculee et la strie mesuree (°).

Site

Type of rock

Age

ol a 2 a 3

ANG

limestone

Late Sarmatian

12409 34977 21609

0.39

limestone

Late Sarmatian

173 27 009 62 066 07

0.38

limestone

Late Sarmatian

351 06 248 65 084 24

0.24

hard volcanic tuff

Badenian

129 04 219 05 003 84

0.58

marl

Oligocene

1 1

29540 20500 11550

0,52

same faults backtilted

1 1

12509 21609 35177

0.52

sandstone

Burdigalian-Langhian

306 01 036 15 210 75

0.57

305 05 056 77 214 12

0.32

285 81 036 03 126 08

0.17

clay

Romanian

148 17 053 16 281 66

0.06

marl

Burdigalian-Langhian

138 50 331 39 236 07

0.18

same faults backtilted

316 08 188 77 047 10

0.18

341 75 076 01 166 15

0.32

216 11 124 11 351 74

0.83

1 1

039 01 307 53 130 37

0.21

limestone

Late Sarmatian

11840 24938 00227

0.24

same faults backtilted

142 03 234 24 047 66

0.24

121 43 006 24 256 37

0.48

same faults backtilted

139.03 038 76 230 14

0.48

1 1

186 06 074 75 277 14

0.48

Northern limb of anticline 9

004 06 119 76 273 13

0.32

Northern limb of anticline 9

220 80 000 08 091 06

0.23

sandstone

Meotian

*07

198 41 094 15 349 45

0.47

same faults backtilted

011 07 279 11 132 78

0.47

1 1

marl

Romanian

190 28 030 60 284 09

0.46

1 1

146 70 027 10 294 18

0.39

clay

Dacian

183 27 075 32 305 46

0.05

054 82 175 04 266 07

0.39

conglomerate

Early Pleistocene

347 13 078 06 193 75

0.33

marl

Meotian

179 06 271 17 071 72

0.64

1 1

clay

Pontian

200 01 297 78 110 12

0.45

1 1

02353 18536 28208

0.05

1 1

clay

Romanian

34107 250 08 113 79

0.54

hard volcanic tuff

Badenian

315 41 054 10 154 47

0.40

same faults backtilted

216 67 058 21 325 08

0.40

319 27 142 60 253 12

0.45

same faults backtilted

032 74 139 05 231 15

0.45

084 23 195 40 332 42

0.63

same faults backtilted

288 09 018 01 117 81

0.63

1 1

134 00 044 02 225 88

0.41

145 78 321 12 051 01

0.42

192 21 331 63 096 16

0.51

limestone

Late Sarmatian

311 01 216 75 041 15

0.14

21176 30902 04014

0.02

limestone

Late Sarmatian

109 06 353 76 201 12

0.38

marl

Burdigalian-Langhian

11331 35439 22835

0.05

limestone

Late Sarmatian

349 34 253 08 152 55

0.26

1 1

same faults backtilted

148 03 238 01 346 87

0.26

sandstone

Badenian

11425 31663 20809

0.38

sandstone

Sarmatian

190 51 069 23 325 30

0.16

same faults backtilted

136 06 226 01 326 84

0.16

Source:

120

JEAN-CLAUDE HIPPOLYTE & MIRCEA SANDULESCU

Fig. 6 .— Cross-section 2. in the Wallachian fold and thrust area (location in Fig. 2; from STEFANESCU, 1984, modified) with

Schmidt diagrams (lower hemisphere) of Wallachian striated fault planes and computed stress axes (bedding planes as

broken lines). Same legend as figure 4.

F/G. 6 — Coupe 2, dans la zone des chevauchements et plis Valaques (localisation en figure 2 ; modifii d'apres STEFANESCU

1984) avec diagrammes (projection Schmidt, hemisphere inferieur) de plans de failles stries d'dge Valaque, axes de

contraintes calcules. Mime legende que figure 4.

confirm the chronology between these two compressional trends. This chronology is also confirmed by

simple geometrical relationships between computed stress axes and bedding planes. For example, the

stress axes of the NW-SE compression are tilted at sites 9 and 17 (table 1 and annex) (the compression

faulted the rocks before bed-tilting), while in the same sites, the stress axes of the N-S compression are

not tilted and therefore postdate bed-tilting.

The NW-SE compression predating Meotian time characterizes the Moldavian 'phase' (SANDULESCU,

1980) when large nappe emplacement occurred. This compression was also responsible for folding. In

site 17, the pre-tilt and post-tilt characteristics of the stress axes of the NW-SE compression indicate that

the folding of this area was completed before the N-S Meotian-early Pleistocene compression. As for the

Wallachian event, during this NW-SE compression the stress regime was either reverse (sites 4, 5, 17,

20, 21, 23, table 1) or strike-slip (sites 1,8, 18, 19, 22, table 1) and temporal stress permutations,

between reverse and strike-slip regimes, also occurred (sites 6 and 9, table 1 and annex).

For both the N-S and the NW-SE compression, the temporal stress permutations could lead to normal

faulting with extension perpendicular to the direction of compression (sites 9B, 11, 12, 15, 17, 18, table

1 and annex). This is a common feature which is no more frequent in this foothill area, than for example

in autochtonous foreland carbonate platforms. One cannot conclude that this extension results from

diapir tectonics, since it is not found above diapirs (MRAZEC, 1927) and, on the contrary, it is sometimes

found in synclines (sites 11, 15, 17). This phenomenon only affects reduced areas of some outcrops. Its

local development within major compressional structures indicates that these stress permutations

probably reflect lateral variations in the confining pressure during the deformation, rather than a general

extensional event.

The two directions of compression are very different (Figs 2, 7). The existence of block rotations

could not be controlled in this area by paleomagnetism because the tested Neogene rocks gave week and

Source: MNHN. Paris

NEOGENE EVOLUTION OF THE CARPATHIAN FOOTHILLS

121

Pioiesti/^ foredeep

Buzau

L j

Fig. 7.— Direction of compression of Moldavian age (early Miocene-Sarmatian). Same legend and location as fgure 2.

FlG. 7. — Direction de compression d'age Moldave (Miocene inferieur-Sarmatien). Meme legende el localisation que pour la

figure 2.

Pig. 8.— Neogene evolution of the Romanian diapir fold area, a) End of the nappe emplacement of the Moldavides controled

by the NW-SE Miocene compression, b) Thrusting and folding in the Moldavides during the Wallachian (late Miocene,

Pliocene and early Quaternary) tectonic phase. The foredeep develops partly above the Moldavian nappes. It is bounded

to the north by an E-W trending folded belt superposed on the arcuate Moldavian thrust pile.

FlG. 8 .— Evolution Neogene de la zone des plis diapir. a) Fin de la mise en place des nappes Moldaves sous Faction d'une

compression miocene orientee NW-SE. b) Chevauchement et plissement des Moldavides pendant la " phase Valaque ”

(Miocene superieur. Pliocene et Pleistocene inferieur). L'avant-fosse se forme en partie an dessus des nappes. Elle est

limitee au nordpar une chaine plissee d'orientation E-W superposee a Fare des nappes moldaves.

Source:

JEAN-CLAUDE HIPPOLYTE & MIRCEA SANDULESCU

122

unstable natural remanent magnetization. The hypothesis that these two different trends of compression

could result from rotations of fractured blocks within a single stress field is unlikely. The chronology of

the two trends of compression (NW-SE, first and N-S second) indicates that the rocks affected by the N-

S compression would have undergone a counter-clockwise rotation during Neogene to obtain the NW-

SE trend However the existing paleomagnetic data in Miocene and older rocks in the south-east

Carpathians (BAZHENOV et al., 1993), South Carpathians (PATRASCU et al ., 1992) and the Apusem

mounts (BORDEA et al.. 1993; PATRASCU et al., 1994) all indicate clockwise rotations.

Moreover, we can consider that the older trend of compression that was characterized (NW Sb

Miocene event) may have undergone the larger rotation. However, this Miocene event is characteiized

with the same trend until central-south Carpathians (RATSCHBACHER et al., 1993). Close to this area, the

Paleogene and Neogene andesites of the South Apuseni Mountains (Fig. 1) allowed a paleomagnetic

control of the block rotations, showing that no rotation occurred in rocks of middle-late Miocene age (15

to 9 My, PATRASCU et al., 1994). In the non-rotated andesites of this area, we characterize a strike-slip

deformation that has also a NW-SE trend of compression, and is dated of Badenian-Sarmatian by the

filling of pull-apart basins nearby. We conclude that the NW-SE Miocene trend of compression was not

rotated after the middle Miocene in the South Apuseni mounts. On the same way, the Miocene

compression that is trending NW-SE in the South Carpathians and the Southeast Carpathians may have

usually undergone only minor rotations after the middle Miocene. As our measurements, in the diapir

fold area, are made mostly in Late Neogene rocks, the possible block rotations may not have change the

stress orientations in a significant way.

In conclusion, despite the second order extension generated by local stress perturbations, the

paleostress analysis confirms that the inner margin of the foredeep basin is compressive. The structures

of this margin are polyphase, however, their present dominant E-W orientation results mainly Irom the

Wallachian N-S compression. The overstep sequence of thrusting and the obliquity of the foredeep

margin with the outer thrust front result from the change in the compression direction from NW-SE to

N-S, during late Sarmatian-Meotian (Tortonian, Fig. 3).

DISCUSSION AND CONCLUSION

The nappe emplacement in the outer Carpathians (Moldavides) lasted during most of Miocene and

ended in middle Sarmatian with the emplacement of the Subcarpathian nappe. In the diapir fold area, the

NW-SE compression affects the Tertiary stratigraphic sequence until the late Sarmatian deposits.

Therefore, it is this compression that generated the nappes of the outer Carpathians (Fig. 8a). Since we

did not found this compression in more recent formations, we conclude that this stress field ended in late

Sarmatian, just after the end of the nappe emplacement.

After the nappe emplacement, the foredeep subsided strongly in the southern areas (Fig. 1). In the

diapir fold area (Fig. 2), the foredeep inner margin was built by the N-S compression (Fig. 8b). On this

margin, probable piggyback syn-compressive basins (as the Valea Lunga Basin) began to form in

Meotian time. The climax of this compression, in the early Pleistocene, has allowed the definition of the

Wallachian phase (STILLE, 1924; DUMITRESCU & SANDULESCU, 1968).

In contrast to the previous NW-SE compression, the N-S Wallachian compression was not

responsible for large nappe emplacement, but created thrust faults with unprecedented vertical

movements (Figs 4,~6) accounting for the trap of sediments in piggyback basins and the thick sediment

accumulations in the foredeep basin (Fig. 6). In the area north of Buzau, where the trend ol the margin

curves to the north, the foredeep margin could be transpressive with a left lateral component of

movement in accommodation of a N-S shortening (Fig. 8b). Further paleostress determinations will aim

at determining if the tilt of the inner margin of the Getic Depression (east-trending non lolded segment

of the foredeep basin. Fig. 1) also results from a N-S compression, coeval to the NNE-SSW compression

identified in Paleogene and Early Neogene rocks of central South Carpathians (RATSCHBACHER et al.,

1993). In the study area, the change of the stress orientation from NW-SE to N-S explains the overstep

sequence of thrusting (Fig. 8b).

A southward movement of the thrust stack could be the reason for the change in the stress field.

During early and middle Miocene the NW-SE compression was related to the eastward escape of the

Tisza intra-Carpathian block (ROYDEN et al., 1982, CSONTOS et al., 1992; LlNZER, 1996) or a foreland

underthrusting (SANDULESCU, 1989). Afterward, a blocking along the NNW-trending part of the East

Source: MNHN. Parts

NEOGENE EVOLUTION OF THE CARPATHIAN FOOTHILLS

123

Carpathians could have produced this southward movement of the thrust stack, in the not yet blocked

southeast Carpathians.

At the front of the East Carpathians bend area, the inner margin of the foredeep basin is thus a

compressive margin formed by N-S compression. Its compressional structures are superposed on

Miocene nappes and are in internal position (Fig. 8b). This location and the correlative overstep

sequence of thrusting result from a change in the stress orientation, from NW-SE during the outer

Carpathian thrusting (early Miocene to Sarmatian) to N-S during Meotian-early Pleistocene (Fig. 8).

During this latter period, folding and thrusting mostly generated vertical movements. This characteristic

of the deformation must be linked to the contemporaneous thick accumulation of sediments in the

foredeep and explain the formation of piggyback basins in the foredeep inner margin.

ACKNOWLEDGEMENTS

This work was supported by the Peri-Tethys Program. We are grateful to F. ROURE and L. R)DOR for

their helpful comments on this paper. We thank H. RON for the paleomagnetism analysis of samples of

Neogene sediments. We also thank D. BADESCU, N. BADESCU, P. CONSTANTIN and M. TURTURANU for

their help during the field work and their constructive discussions of the results.

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NEOGENE EVOLUTION OF THE CARPATHIAN FOOTHILLS

125

SEC3

SEC6

SEC9

SEC5

SEC6

secq

SEC 1

SEC9

Source: MNHN. Paris

126

JEAN-CLAUDE HIPPOLYTE & MIRCEA SANDULESCU

SEC9B SEC9B

SEC lO

SEC 1 O

SEC 15

SEC 1 7

-vS'-

Source: MNHN. Pans

NEOGENE EVOLUTION OF THE CARPATHIAN FOOTHILLS

127

SEC 17 SEC 17 SEC 18 SEC1Q

Annex: Schmidt diagrams (lower hemisphere) of striated fault planes and computed stress axes (references in table 1; same

legend as for figure 4). Note that in the cases where faulting occurred before folding, the fractures are presented in two

diagrams: one for the present fault attitude (the bedding planes, the faults and the stress axes are tilted) and one for the

original fault attitude (the bedding planes and two stress axes are horizontal).

Annexe : Diagrammes (projection Schmidt, hemisphere inferieur) de plans de failles et d’axes de contraintes calcules

(references dans le tableau / ; me me legende que pour la figure 4). Remarquer que dans les cas ou la fracturation a

precede un plissement important, les failles sont representees dans deux diagrammes : un pour la geometric actuelle

(plans stratigraphiques, failles et axes de contraintes bascules) et un pour la geometric originelle (plans de

stratification et deux axes de contraintes sont horizontaux).

The Moesian Platform as a key for understanding

the geodynamical evolution of the

Carpatho-Balkan alpine system

Frangoise BERGERAT" 1 , Pierre MARTIN' 2 ' & Dimo DlMOV ,sl

Departement de Geotectonique, CNRS URA 1759, case 129, Universite P. et M. Curie

75252 Paris Cedex 05, France

<2 ’ SGN/12G/GEO, B.R.G.M.. B.P. 6009, 45060 Orleans Cedex 2, France

Geology and Geography Faculty, St-Kliment Ohridski University, 15 Tzar Osvoboditel Bd.

1000 Sofia, Bulgaria

ABSTRACT

From the Mesozoic to Present, Moesia has been a foreland platform regardless of the configuration of the Alpine orogen.

The analysis of brittle structures characterizes the successive paleostress fields related to the different Alpine tectonic events of

the Balkan area. Several main fault systems have been identified. A rather extensional phase, N-S to NNE-SSW, characterized

by strike-slip and normal faulting, is of post-Aptian and pre-Maastrichtian age. A N-S to NNE-SSW compressional stress,

characterized by strike-slip and reverse faulting, corresponds to a tectonic phase at the end of the middle Eocene. Another NNE-

SSW compressional stress is of Pliocene age. An E-W extensional stress may by related to the N-S compression of middle

Eocene age (permutation of stresses) or to a Miocene phase. The question of possible rotation and / or translation of the

Moesian platform is discussed taking into account paleomagnetic data. A major rotation has not occurred. No data can confirm

or invalidate a westward translation of Moesia. The main identified tectonic events are connected both to the Black Sea opening

and to the Eurasian and Arabo-African plates convergence.

RESUME

La plate-forme moesienne, cle pour la comprehension de revolution geodynamique alpine du systeme Carpates-

Balkans.

Du Mesozoique a l'Actuel. la Mo6sie a constitue une plate-forme d'avantpays de 1'orogene alpin. L'analyse des structures

cassantes caracterise les paleo-champs de contrainte successifs lies aux differents evenements techniques alpins des Balkans.

Plusieurs systemes de failles majeurs ontete identifies. Une phase a dominante extensive. N-S a NNE-SSW, caracterisee par

des decrochements et des failles normales, est d'age post-Aptien et ante-Maastrichtien. Une compression N-S a NNE-SSW,

caracterisee par des decrochements et des failles inverses, correspond a une phase technique datee de la fin de l'Eocene moyen.

Une autre compression NNE-SSW est d'age pliocene. Une contrainte extensive E-W peut etre liee soit a la compression N-S de

l’Eocene moyen, par un phenomene de permutation de contraintes, soit a une phase miocene. La question de la rotation et / ou

de la translation de la Moesie est discutee en prenant en compte les donnees du paleomagnetisme. Aucune rotation majeure n'a

pu se produire. Aucune donnee ne permet de confirmer ou d'infirmer une translation vers l'ouest de la plate-forme moesienne.

Les mouvements techniques majeurs identifies sont attribues a la fois a l'ouverture de laMer Noire et h la convergence Afrique

/ Eurasie.

Bergerat, F, Martin. P. & Dimov. D., 1998.— The Moesian Platform as a key for understanding the geodynamical

evolution of the Carpatho-Balkan alpine system. In: S. CrasqUIN-Sdleau & E. Barrier (eds), Peri-Tethys Memoir 3:

stratigraphy and evolution of Peri-Tcthyan platforms. Mem. Mus. natn. Hist, nat 177 : 129-150. Paris ISBN : 2-85653-512-7.

Source: MNHN, Paris

130

FRANCOISE BERGERAT ETAL.

INTRODUCTION

Various geological synthesis have been proposed for Tethyan and peri-Tethyan chains, platforms and

basins (e.g. DERCOURT et a /., 1985, 1986; RlCOU et al ., 1985, 1986). In these syntheses, the position

occupied by the Moesian platform brings up several questions regarding its relationships with the

neighboring chains of the Carpathian Mountains and the Balkan (Fig. 1). One of the main questions

concerns its possible rotation or translation relative to stable Europe.

The present Moesian platform is part of a foreland of the Carpatho-Balkan chain. In Romania, it

consists essentially of the Carpathian fore-deep, where Pliocene and Miocene molassic formations are

largely developed. In Bulgaria, it is represented in its central and western parts by Cretaceous and

Paleocene formations that have mostly subhorizontal bedding dips. The Miocene and Pliocene

formations crop out to the North-West and in the East of the platform in two depressions that have

formed since the middle Miocene.

Fig. I.— Tectonic sketch map of the Central European Alpine System. Abbreviations are as follows: A, Apuseni Mountains;

EA, Eastern Alps; EC, Eastern Carpathians; IMF. Intra-Moesian Fault; MF, Maritsa Fault; ND, Northern Dobrogea; PB,

Po Basin; PCF, Peceneaga-Camena Fault; PK, Pieniny Klippen zone (hatched area), SC, Southern Carpathians; TB,

Transylvanian Basin; WC. Western Carpathians.

FlG. /.— Carte technique schematique du systeme alpin d’Europe Centrale. A, Monts Apuseni; EA, Alpes orientates ; EC,

Carpates orientates ; IMF, faille intra-moesienne ; MF, faille de la Maritsa ; ND, Dobrogee-Nord ; PB : bassin du Po ;

PCF, faille de Peceneaga-Camena ; PK, zone des klippes piennines (hachures) ; SC, Carpates meridionales ; TB,

bassin transylvain ; WC, Carpates occidentales.

Source: MNHN. Paris

THE MOESIAN PLATFORM: GEODYNAMICAL EVOLUTION

131

Location of microtectonic measurements sites (stars, 1 to 50) and samples collection for paleomagnetic studies (squares,

1 to 9). Numbers refer to the following: 1, Plio-Quaternary; 2, Plio-Pleistocene basalts; 3, Mio-Pliocene of Lorn and

Vidin depressions and of the Carpathian foredeep; 4 , Cenozoic intra-Balkan basins; 5 , Cretaceous of the Moesian

platform; 6 , Fore-Balkan and Stara Planina; 7 , Sredna Gora; 8 . Rhodope-Strandza; 9 . Precambrian and Paleozoic of

Dobrogea; 10 , major thrusts.

Abbreviations are as follows: Ba. Balcik; Be, Bercovica, Bk. Belogradcik; Bu, Burgas; D, Dragoman; E. Etropole; J,

Jablanica; K, Kula; L, Lorn; N, Nis; Pr, Preslav; PL, Plovdiv; R, Russe; S, Svoge; Si, Silistra; T. Trojan; Va, Varna; Vi,

Vidin; VT, Veliko Tamovo; TF, Timok Fault.

FlG. 2— Carte geologique simplifiee de la plate-forme moesienne et du domaine balkanique (d’apres la carte geologique

internationale de I'Europe a l/l 500 000, 1969-1983, et les cartes geologiques de Bulgarie a 1/500 000, 1989 et

1/1 000 000, 1978).

La localisation des sites de mesures microtectoniques est indiquee par des etoiles (numerotees de 1 a 50), celle des

echantillons pour les etudes de paleomagnetisme par des carres (numerotes de 1 a 9). 1, Plio-Quaternaire ; 2, basaltes

plio-pleistocenes ;3, Mio-Pliocene des depressions de Lorn et de Vidin et de I'avant-fosse carpathique ; 4 . bass ins

cenozoi'ques intra-balkaniques ; 5. Cretace de la plate-forme moesienne ; 6, Prebalkan et Stara Planina ; 7, Sredna-

Gora ; 8, Rhodope-Strandza ; 9, Precambrien et Paleozotque de la Dobrogee ; 10, chevauchements majeurs.

The classic division of Bulgaria was established by BONCEV (1940) and is still used (e.g. BONCEV,

1988; IVANOV, 1988). South of the Moesian platform and the small triangle of Kula attributed to the

Southern Carpathians (Fig. 2), three main zones occur, the Fore-Balkan, the Stara-Planina (including the

subzone of Luda Kamcija) and the Sredna Gora. The Rhodope is located to the South of the Sredna

Gora. However, these various zones and subzones correspond to superposed alpine configurations

(DERCOURT & RlCOU, 1987). Therefore these names will only represent tectonic zones here if an age is

associated. In the opposite case, they will have only a meaning of geographical sector.

The main tectonic and paleogeographical configurations, from the Jurassic to the Tertiary, are clearly

oblique to one another. Furthermore, several of these paleogeographical elements have played a major

role in the Alpine development of Bulgaria. The basin of Nis-Trojan of Upper Jurassic-Lower

Cretaceous age, and the magmatic axis of the Upper Cretaceous are included with these elements

(DERCOURT & RlCOU, 1987). These major structures have influenced later deformations. The inherited

structure and the obliquity of successive tectonic configurations, make the structure of the Balkan

extremely complex.

In northern Bulgaria, the zone corresponding presently to the Moesian platform has always

132

FRANCOISE BERGERAT ETAL.

represented a foreland platform in the different tectonic settings from the Mesozoic to the Present. The

analysis of brittle deformations, can be used to characterize the paleostress fields associated to the

different Alpine tectonic events.

The aims of this paper are a- to describe and interpret in terms of paleostress tensors the different

fault systems that affect the Cretaceous to Pliocene formations of the Moesian platform, b- to examine

the possible rotation and / or translation of this platform, and c- to discuss the paleostresses deduced in

the framework of Black Sea opening and Africa-Eurasia convergence.

ALPINE DEVELOPMENT OF THE MOESIA-BALKAN DOMAIN

In the Triassic, a continental platform developed on the post-Hercynian peneplain. The same

subsiding platform facies can be observed from the foreland to the Central Sredna Gora.

In the latest Jurassic-Lower Cretaceous, the Nis-Trojan flysch basin (3000 m of Lower Cretaceous)

developed at the south-eastern margin of the platform (Fig. 3). It separates the Moesian platform to the

North from a series of horsts to the South. The Jablanica line (NlCOLOV & KRISHEV, 1965), also called

the Etropole line, represents zone of variable width of facies changing. For DERCOURT & RICOU (1987),

this line is a transform fault linked to the opening of the basin. IVANOV (pers. comm.) believes the basin

is bounded to the SSE by a strike-slip fault, separating it from fragments of the Serbo-Macedonian

massif. The basin disappears in the Aptian. During the Albian, an important tectonic phase is

characterized by folds approximately E-W and thrust faults verging North (Ivanov, 1988). This tectonic

event appears to copy almost exactly the previous setting because the deformed zone front corresponds

approximately to the ancient Nis-Trojan basin (DERCOURT & RlCOU, 1987).

Fig. 3.— Superimposed tectonic and paleogeographic settings in the Moesian platform and Balkan domain.

Numbers refer to the following features: 1, Eo-Cretaceous Moesian platform (including Fore-Balkan); 2, Nis-Trojan

basin, Sevenn (S) and Kraina-Kula Klippes; 3 , Neo-Cretaceous Getic (G) domain , and of Zemen (Z) and Sredna Gora

vl' • Luti o^ aiI| c |j a basin; 5 ' Neo-Cretaceous magmatic axis; 6, Rhodope-Strandza; 7, Cenozoic intra-Balkan basins,

Krawte, SB, South-Balkan, UT, Upper Thrace; 8, major Eocene thrusts; 9, major Miocene strike-slip faults

Abbreviations for towns as for figure 2.

F,G - 3 -— Techniques et paleogeographies superposees dans la plate-forme moesienne et le domaine balkanique.

1 plate-forme moesienne eo-cretacee (incluant le Prebalkan) ; 2. bassin de Nis-Trojan, Klippes de Severin (S) et

Krama-Kula ; 3, domaines neo-cretaces getique (G). de Zemen (Z) et de Sredna Gora (SG) ; 4. bassin de Ltida

Kamcija; 5, axe magmatique neo-cretace : 6, Rhodope-Strandza ; 7. bassins cenozoiques intra-balkaniques ; 8,

chevauchements eocenes majeurs ; 9, decrochements miocenes majeurs.

Source: MNHN. Paris

THE MOESIAN PLATFORM: GEODYNAMICAL EVOLUTION

133

Starting during the Senonian, a new flysch basin developed (NACHEV, 1977, 1978, 1980) in the

Eastern and Northern part of the Stara Planina (Fig. 3). It constitutes the subzone of Luda Kamcija

(flysch of Emine of Turonian-lower Paleocene age). During the same period to the South of this zone,

significant calc-alkaline magmatism, active especially during the Coniacian-Santonian (more than 6000

m of volcano-clastic sediments), occurs in the Srednogorian rift (Fig. 3). The volume of volcanogenic

rocks and their potassium content increase towards the Black Sea, but the neocretaceous magmatic axis

can be recognized throughout Bulgaria, from Dragoman to Burgas (Fig. 3). This axis is superposed

obliquely upon the ancient structures. Furthermore, the non-volcanic facies (flysch of Emine) are known

only to the North of the axis and to the East of the Etropole line. Elsewhere in the Stara Planina as well

as in the Fore-Balkan, a platform style sedimentation prevailed.

During the lower and middle Eocene, a vast molassic depression, trending NW-SE to E-W, extending

from Serbia to the Black Sea (NACHEV, 1980) developed. These molasse deposits have been folded after

the middle Eocene. The formation of folds in the Stara Planina and the Fore-Balkan has been

immediately followed by significant thrusting to the North (Fig. 3) (IVANOV, 1988). These folds and

thrusts characterize another major tectonic phase. The most significant thrust is the so-called Stara

Planina thrust (thrusting of the Sredna Gora over the Stara Planina). Some other thrusts also

accommodate the shortening on the northern flanks of the Svoge, Bercovica and Belogradcik anticlines

(Fig-3)- J

From Neogene to Quaternary, the post-tectonic configuration superposed on the various alpine

elements. Several systems of troughs exist (Fig. 3). a- The kraistides trend N 150°-160° and separate

Serbo-Macedonian and Rhodope, b- The South-Balkan is oriented E-W and occurs between Stara

Planina and Sredna Gora. c- The depression of Upper Thrace has an irregular form and masks the

boundary between Sredna Gora and Rhodope.

BRITTLE DEFORMATION AND PALEOSTRESSES IN THE MOESIAN PLATFORM

Except for domes and depressions, the deformation of the Moesian platform is essentially

characterized by brittle tectonics (BERGERAT & PlRONKCTv, 1994). The most ancient exposed rocks

suitable for tectonic analysis are of Lower Cretaceous age and the most recent are of Pliocene age. The

brittle deformation study has been carried out in 50 sites of which 22 are in calcareous formations of

Lower Cretaceous age (Valanginian to Aptian), 9 are in Maastrichtian-Paleocene calcareous formations,

12 are in Miocene formations and 7 are in the Pliocene. All sites of measurements were located in the

platform itself, except 1 in the Fore-Balkan (Fig. 2).

Joints, tension gashes, tectonic stylolites and faults all occur at 45 sites. Most sites contain evidence

of several tectonic events. The various populations of faults were separated by using the mechanical

coherence criteria and chronological observations in the field (BERGERAT, 1994) such as crosscutting

structures, tilting of early-formed structures and superposition of slickenside lineations on the same

plane (Figs 4, 5). This analysis allowed us to distinguish 5 main fault systems (including 2 geometrically

identical, but of different ages) that have been analyzed in terms of paleostresses (INVDIR method;

ANGELIER, 1990).

E- W TO ESE- WNW COMPRESSION AND N-S TO NNE-SSW EXTENSION

A right-lateral strike-slip system with an azimuth 40° to 60° and a left-lateral one with an azimuth

150° to 160° have been identified in 25% of the measured sites in the Moesian platform (Fig. 6). These

sites are all located in limestones of Valanginian to Aptian age with subhorizontal bedding dips. This

fault system has not been observed in formations of Upper Cretaceous age or more recent. Moreover

chronological criteria commonly indicate that it is prior to the approximately N-S compression and / or

E-W extension.

Only 5 sites present a sufficient number of faults to obtain a stress tensor of good quality. The

azimuth of the maximal principal stress ol varies from 95° to 120° and is horizontal (Table 1).

134

FRANCHISE BERGERAT ET AL.

Fig. 4.— Analysis of the Carev Brod quarry (site 4, Hauterivian-Valanginian limestones), a polyphased site (after Bergerat &

Pironkov, 1994). Separation of the different stress states and relative chronological criteria.

I. Entire fault-slip data population represented on stereographic projections of fault planes (great circles) and slickenside

lineations (divergent arrows for normal slip, convergent arrows for reverse slip, double arrows for strike-slip). Data are

plotted on lower hemisphere, equal-area projections.

II. Separation of the entire population into three sets of stress tensors (cf. tables), a. Strike-slip faults compatible with

ESE-WNW compression/NNE-SSW extension, b. Strike-slip faults compatible with a NNE-SSW compression/ESE-

WNW extension, c. Normal faults compatible with a E-W extension. Five, four and three-prong stars represent 0 1, (72

and (73 respectively. Large black arrows indicate inferred directions for compression and extension.

III. Examples of fault planes bearing two generations of slickenside lineations. A: NNW-SSE trending fault plane with

early horizontal striations showing left-lateral motion (set a) and a second one showing right-lateral motion (set b); B:

Another fault plane with the same trend showing early left-lateral horizontal striations (set a) and a later dip-slip one

showing normal motion (set c).

Fig. 4.— Exemple de site polyphase : la carriere de Carev Brod (site 4, calcaires de I'Hauterivien-Valanginien) (d'apres

Bergerat & Pironkov, 1994). Separation des different etats de contrainte et criteres de chronologie relative.

I. Lot de donnees brutes : projections cyclographiques des plans de failles et de leurs stries (canevas de Schmidt,

hemisphere inferieur, symbolise sur les diagrammes par un S dans un demi-cercle), fleches centrifuges : failles

normales, fleches centripetes : failles inverses, doubles fleches : decrochements.

II. Separation de la population entiere de failles en trois families. Analyse en termes de tenseurs des contraintes (cf.

tableaux), a, decrochements compatibles avec une compression ESE-WNW et une extension NNE-SSW. b,

decrochements compatibles avec une compression NNE-SSW et une extension ESE-WNW, c, failles normales

compatibles avec une extension E-W. Etoiles a 5, 4 et 3 branches : Cl. (72 et C3 respectivement, grosses fleches

noires : directions probables de compression et dextension.

III. Exemples de plans de failles portant deux generations de stries, A . plan de direction NNW-SSE portant une

premiere striation, horizontale, caracteristique d'un jeu senestre (famille a) et une seconde egalement horizontal,

posterieure, caracteristique d'un jeu dextre (famille b) ; B : autre plan de me me direction portant egalement une

premiere striation horizontale montrant un jeu senestre (famille a) et une seconde " dip-slip " indiquant un mouvement

normal (famille cj.

Source: MNHN. Paris

THE MOESIAN PLATFORM: GEODYNAMICAL EVOLUTION

135

i nr

Fig. 5 .— Example of polyphased site at roadcut near Devnja (site 43, Hauterivian-Valanginian limestones). Separation of the

different stress states and relative chronological criteria.

I. Entire fault-slip data population.

II. Separation of the entire population into three sets. Analysis done in terms of stress tensors (cf. tables), a. Strike-slip

faults compatible with a ESE-WNW compression/NNE-SSW extension, b. Strike-slip faults compatible with a N-S

compression/E-W extension, c, Normal faults compatible with an E-W extension.

III. Examples of fault planes bearing two generations of slickenside lineations. A: NW-SE trending fault plane bearing

early horizontal striations showing left-lateral motion (set a) and a second one showing a right lateral motion (set b). B:

A NNW-SSE trending fault plane showing early left-lateral horizontal striations (set a) and dip-slip striations showing a

normal motion (set c).

Fig. 5.— Exemple de site polyphase : bord de route pres de Devnja (site 43, calcaires de I'Hauterivien-Valanginien).

Separation des differents etats de contrainte el criteres de chronologic relative.

I. Lot de donnees brutes

II. Separation de la population entire de failles en trois families. Analyse en termes de tenseur des contraintes (cf.

tableaux), a, decrochements compatibles avec une compression ESE-WNW et une extension NNE-SSW,

b, decrochements compatibles avec une compression N-S et une extension E-W, c, failles normales compatibles avec

une extension E-W.

III. Exemples de plans de failles portant deux generations de stries. A : plan de direction NW-SE portant une premiere

stnation, horizontale, caracteristique d'un jeu senestre (famille a) et une seconde, egalement horizontal } e, posterieure,

caracteristique d'un jeu dextre (famille b). B : plan de direction NNW-SSE portant une premiere striation, horizontale,

montrant un jeu senestre (famille a) et une seconde ''dip-slip'' indiquant un mouvement normal (famille c).

NNW-SSE TO NNE-SSW COMPRESSION AND ENE-WSW TO WNW-ESE EXTENSION

In 50% of the sites of the Moesian platform many strike-slip faults give a state of stress with

horizontal compressive stress axis NNW-SSE to NNE-SSW. Many horizontal stylolitic peacks, trending

NNE-SSW. accompany commonly the strike-slip system (Fig. 7). Identified in all formations of

136

FRANCOISE BERGERAT ETAL.

Fig. 6 .— ESE-WNW to E-W compression in the Moesian platform.

Large black arrows show computed directions of compression, small black arrows show estimated directions of

compression.

Fig. 6 .— Compression ESE-WNW a E-W dans la plate-forme moesienne.

Us grandes fitches noires indiquent les directions de compression calculees, les petites fitches noires, les directions de

compressions estimees.

N°

Locality

Age

a 3

RUP

5-6

CAREV BROD

Valanginian-Hauterivian

120-8

250-78

28-7

0.7

GARA HITRINO 1

Valanginian-Hauterivian

94-5

285-85

184-1

0.4

TOPCII

Aptian

272-10

80-80

181-2

0.2

ZLATNA

Berriasian-Barremian

306-15

123-75

215-1

0.4

DEVNJA

Valanginian-Hauterivian

94-48

297-39

197-12

0.7

Table 1Paleostress tensor computations based on fault slip data analysis for the ESE-WNW to E-W compression.

The columns contain from left to right: Site number on Figure 2 (N°) and name of the closest locality; age of the rocks;

number of fault slip data (N); trend and plunge of the principal stress axesGl, G2, G3 (in degrees); ratio <P = (G2-G3) /

n, ."? 3): r a u er x?fx?/lfi e bel Y een com P uted sh ear stress and observed slickenside lineation (a in degrees); ratio "upsilon"

(RUP) of the INVDIR method (cf. Angelier. 1990). Possible values of RUP range from 0 to 200%. Average RUP

values below 50% correspond to good fits between actual fault slip data distribution and computed shear stress

distribution. r

Tableau 1. — Resultats des calculs des tenseurs de paleo-contraintes pour la compression ESE-WNW a E-W. De gauche a

d f ro !! e \ num f ro du site (N°, voir localisation figure 2) et nom de la localite la plus proche ; age des roches. nombre de

James a stries (N) ; directions et prolongements des axes de paleo-contraintes Gl, G2 et 03 (en degres) ; rapport des

contraintes principalesQ - { 02 - 03)/( Gl - G3) ; angle moyen entre strie calculee et strie reelle (a en degres) ;

mfiL rt r? P5ll T RUP) d€ U , ! , ^ ode INVDIR ^f. ANGEUER, 1990). Valeurs possibles de RUP comprises entre 0 et

zUU/c. Des valeurs moyennes de RUP inferieures a 50% correspondent a une bonne coherence entre la strie reelle

observee et la contrainte tangentielle calculee.

Source: MNHN. Paris

THE MOESIAN PLATFORM: GEODYNAMICAL EVOLUTION

137

Fig. 7.— NNW-SSE to NNE-SSW compression in the Moesian platform.

Arrows as for Figure 6.

Fig. 7.— Compression NNW-SSE a NNE-SSW dans la plate-forme moesienne.

Legende : voir figure 6. _

Valanginian to Paleocene age, this state of stress is represented by a right-lateral fault system with an

azimuth 160°-200° and a left-lateral one with an azimuth 50°-90°. Near the Fore-Balkan as well as in the

Preslav anticline (Fore-Balkan, Fig. 2), these strike-slip faults are associated with many reverse faults

with an azimuth 90°-135°. The inferred directions of ol are practically identical if we separately

compute the stress tensor for strike-slip faults and for reverse faults, except for a stress permutation of

o2 and o3 (Fig. 8). In detail at the Fore-Balkan site, strike-slip faults are clearly tilted with the layering

and predate the reverse faults indicating an evolution from a strike-slip regime to a purely compressive

regime during the same compression phase.

Fig. 8 .— Example of 02/03 stress permutations at Preslav site.

The entire fault population has been back-tilted around an average bedding plane N 115°-40°S. This fault slip data

population characterizing a NNE-SSW compression (a) may be divided in two sets with (b) strike-slip faulting (Ol and

03 horizontal) and (c) reverse faulting (Ol and 02 horizontal).

Fig. 8.— Exemples de permutations des contraintes 02 et 03 dans le site de Preslav.

La population entiere de failles (a) peut etre divisee en deux families : decrochements caracterisant une compression

NNE-SSW et une extension ESE-WNW. avec Ol et 03 horizontaux (b), et failles inverses caracterisant une compression

NNE-SSW, avec Ol et 02 horizontaux (c).

138

FRANCHISE BERGERAT ETAL.

Table 2._Paleostress tensor computations based on fault slip data analysis for the NNW-SSE to NNE-SSW compression. See

caption of

U 2.— Re Si

tableau 1.

caption of table 1 for definitions. , , , ,

Tableau 2.— Resultats des calculs de tenseurs de paleo-contraintes pour la compression NNW-SSb a NNb-^w. Legenae voir

N°

Locality

Age

o 2

G 3

RUP

PRESLAV

Valanginian-Hauterivian

014-12

105-4

214-77

0.3

019-14

140-64

283-21

0.2

KASPICAN

Valanginian-Hauterivian

33-12

228-77

123-3

0.6

5-6

CAREV BROD

Valanginian-Hauterivian

16-40

188-50

283-4

0.7

GARA HITRINO 1

Valanginian-Hauterivian

211-10

37-79

301-1

0.5

RAZGRAD

Barremian

33-13

177-74

301-9

0.4

PLEVEN

Maastrichtian-Paleocene

351-1

229-87

81-2

0.4

RALOVO

Maastrichtian

184-17

351-73

93-4

0.7

TODOROVO

Maastrichtian

196-3

354-87

106-1

0.4

BEJANOVO

Maastrichtian

15-9

106-13

252-74

0.4

21-22

BOGOROVO

Aptian

18-14

236-72

110-10

0.1

10-24

SREBARNA

Pliocene

27-1

118-65

296-25

0.2

MERKOVO

Danian

32-8

194-82

302-2

0.6

11-21

119-39

260-44

0.2

ZLATNA

Berriasian-Barremian

341-2

244-72

72-18

0.3

210-14

5-75

118-6

0.9

DEVNJA

Valanginian-Hauterivian

359-3

131-86

269-3

0.5

PROF. ICHIRKOVO

Pliocene

34-11

186-78

303-6

0.5

Fig. 9.— Azimuthal dispersion of the maximal principal stress for the N-S to

NNE-SSW compression (18 stress tensors computed).

FtG. 9.— Dispersion azimulale de la contrainte principale maximale pour la

compression N-S a NNE-SSW (18 tenseurs calcules).

Source: MNHN. Paris

THE MOESIAN PLATFORM: GEODYNAMICAL EVOLUTION

139

The direction of ol varies in all the Cretaceous-Paleocene sites of the Moesian platform between

NNW-SSE and NNE-SSW with two maxima at N 15° and N 30° (Fig. 9, Table 2). However, the data do

not support the presence of two distinct tectonic episodes.

A NNE-SSW compressional state of stress has also been recognized in two sites located in the

Pliocene limestones of the region of Silistra (North-East of the Moesian platform; Fig. 2). Despite the

poor quality of these limestones (terrestrial facies unfavorable to a good preservation of the movement

indicators), one can observe left-lateral strike-slip faults of azimuth 50°-70° and right-lateral ones of

azimuth 5°-10°, as well as horizontal reverse faults of azimuth 115°-120° (Fig. 10). These data allow

one to define a horizontal maximum principal stress al with an azimuth 30° (Table 2). A disrupted fold

with an E-W trending axis in Sarmatian limestones near Balcik (site 37, Fig. 2) also shows this

compression.

Fig. 10.— NNE-SW compression in the Pliocene, Northeastern part of the Moesian platform.

a, Reverse faults at Professor Ichirkovo; b. Stereographic projection of faults planes at Professor Ichirkovo roadcut; c.

Stereographic projection of faults planes at Srebama roadcut.

FlG. 10.— Compression NNE-SSW dans le Pliocene de la partie nord-est de la plate-forme moesienne.

a. failles inverses d Professor Ichirkovo ; projections stereographiques des plans de failles et contraintes principales a

Professor Ichirkovo (b) et a Srebama (c).

140

FRANCOISE BERGERAT ETAL.

NNW-SSE TO NNE-SSW EXTENSION

Many normal faults with an azimuth 70° to 130° have been observed in 30% of the sites, exclusively

in rocks of Valanginian to Aptian age (Fig. 11) and characterize an approximately E-W a3 axis. These

normal faults predate the approximately N-S strike-slip compression, as well as the E-W extension. On

the other hand, no chronological relationship has been found between the E-W compression and the

approximately N-S extension.

The direction of extension o3 resulting from the computation of stress tensors varies from N 160° to

N 200° following the dominance of the set of faults with an azimuth 70°-80° or that with an azimuth

120°-130° (Table 3). There is currently no evidence to support 2 distinct fracture sets. Given the lack of

chronological relations between the N-S compression and the E-W extension and the general

compatibility of these two stress regimes, it is likely that they represent a single phase with a

permutation of o2 and a3 (Fig. 12).

Fig. 11.— NNW-SSE to NNE-SSW extension in the Moesian platform.

Large black arrows show computed directions of extension, small black arrows show estimated directions of extension.

Fig. II .— Extension NNW-SSE a NNE-SSW dans la plate-forme moesienne.

Les grandes fleches noires indiquent les directions d'extension calculees, les petites fleches noires, les directions

d'extension estirnees.

E-W TO WNW-ESE EXTENSION

Normal faults with an azimuth 170°-200° are present in 49% of the sites of Valanginian to Paleocene

age and allow the characterization of a state of extensional stress E-W to WNW-ESE (Fig. 13). In most

sites, the direction of extension is close to N100°-l 10°, with extreme values of N80° and N135° (Table

4).

The systematic presence of both strike-slip and normal faults corresponding to a same direction of o3

in most of the sites (Table 1), as well as the systematic absence of chronological criteria between these 2

fracture sets, suggest that the difference between these extensional and strike-slip regimes is a

permutation of stress o2 and g 3 (Fig. 12).

Nevertheless one station in the Sarmatian of the region of Vidin (North-West of the Moesian

platform; Fig. 2) also recorded an E-W trending extension (Fig. 14) that is necessarily younger than the

Source: MNHN. Paris

THE MOESIAN PLATFORM: GEODYNAMICAL EVOLUTION

141

Table 3.— Paleostress lensor computations based on fault slip data analysis for the NNW-SSE to NNE-SSW extension. See

caption of table 1 for definitions.

Tableau 3 .— Resultats des calculs de tenseurs de paleo-contraintes pour Textension NNW-SSE a NNE-SSW . Legende voir

tableau 1.

N°

Locality

Age

a 1

RUP

TOPCII

Aptian

354-71

114-10

207-16

0.2

CERVEN

Aptian

41-71

289-8

196-17

0.5

NICOVO

Aptian

358-72

256-4

164-17

0.3

TITOVO

Aptian

277-81

76-9

166-3

0.2

TERTER

Aptian

39-85

249-5

159-3

0.2

21-22

BOGOROVO

Aptian

310-77

183-8

91-10

0.0

GLAVINICA

Aptian

23-74

277-5

186-16

0.3

STOJAN MIHAJ.

Berriasian-Barremian

133-85

260-3

351-4

0.4

Fig. 12.— Examples of G 1/<J2 permutations, sites of Topcii (A) and Kaspican (B).

In each site, the entire fault slip data population characterizing NNE-SSW extension (A. a) or ESE-WNW extension (B.

a) may be divided into two sets with (b) strike-slip faulting (Gl and 03 horizontal) and (c) normal faulting (G2 and 03

horizontal).

FlG. 12 .— Exemples de permutations des contraintes Ol et 02 : sites de Topcii (A) et Kaspican (B).

Dans chaque site la population entiere de failles caracterisant me extension NNE-SSW (Aa) ou ESE-WNW (Ba) peut-

etre divisee en deux families : des decrochements avec Gl et G3 horizontaux (b) et des failles normales, avec G2 et 03

horizontaux (c).

Source

142

FRANCHISE BERGERAT ETAL.

Fig. 13.— E-W 10 WNW-ESE extension in the Moesian platform.

Arrows as for figure 11.

FlG. 13 .— Extension E-W a WNW-ESE dans la plate-forme moesienne.

Legende : voir figure 11.

Fig. 14.— E-W extension in the middle Miocene at Dimovo.

a- Normal faults in tilted sandstone layers (average bedding plane N 180°-20°E), b- Stereographic projection of fault

planes after back-tilting. White arrows as estimated direction of extension.

Fig. 14 .— Extension E-W dans le Miocene moyen a Dimovo.

a . failles normales dans les couches greseuses basculees (plan de stratification moyen : N 180°-20°E), h : projections

stereographiques des plans de failles apres remise a Vhorizontale de la stratification. Les fleches blanches indiquent la

direction d'extension estimee.

Source: MNHN. Paris

THE MOESIAN PLATFORM: GEODYNAMICAL EVOLUTION

143

Table 4.— Paleostress tensor computations based on fault slip data analysis for the E-W to ESE-WNW extension. See caption

of table 1 for definitions.

Tableau 4 .— Resullats des calculs de tenseurs de paleo-contraintes pour lextension a ESE-WNW. Legende voir tableau 1.

N°

Locality

Age

<x>

RUP

KASPICAN

Valanginian-Hautcrivian

263-83

25-4

115-6

0.2

MATNICA

Valanginian-Hauterivian

264-77

5-3

96-13

0.1

SUMEN

Senonian

276-82

17-2

107-8

0.4

5-7

CAREV BROD

Valanginian-Hauterivian

100-81

2-1

272-9

0.4

GARA H1TRINO 1

Valanginian-Hauterivian

287-71

42-8

135-17

0.2

RAZGRAD

Barremian

160-87

20-2

290-2

0.5

TOPCII

Aptian

287-78

195-1

104-12

0.2

TETOVO

Aptian

226-73

354-10

86-13

0.2

TERTER

Aptian

85-73

188-4

279-17

0.4

21-22

BOGOROVO

Aptian

359-82

197-8

107-3

0.2

BELICA

Aptian

45-77

168-7

259-11

0.3

GARA HITRINO 2

Valanginian-Hauterivian

111-78

-209-2

299-12

0.3

TO JAN MIHAJ.

Berriasian-Barremian

77-83

182-2

272-7

0.3

ZLATNA

Berriasian-Barremian

272-84

7-1

97-6

0.5

DEVNJA

Valanginian-Hauterivian

355-65

192-24

99-7

0.3

BACARBOVO

Berriasian-Barremian

108-74

199-0

289-16

0.4

deformation discussed above. Obviously we cannot assert from a single site located in the northwest

extremity of the platform that this deformation represents the same tectonic phase because no normal

faults have been observed in the Miocene of the Balcik-Vama area.

DISCUSSION

The strike-slip stress state characterized by a horizontal maximal principal stress al E-W to ESE-

WNW has never been identified in rocks younger than the Aptian. Moreover, in sites of Valanginian to

Aptian age where it could be demonstrated, relative chronology criteria (Fig. 4) indicate that it is prior to

states of stress in compression approximately N-S and / or extension approximately E-W, which affect

not only the Lower Cretaceous, but also the Upper Cretaceous-Paleocene.

Similar remarks may be made concerning the state of stress characterized by normal faults and a

NNW-SSE to NNE-SSW horizontal principal minimal stress o3. This one is also only observable in

rocks of Valanginian to Aptian age and always predates the N-S compression and the E-W extension.

The age and directions of o3 are identical for these two states of stress implying that it concerns a single

phase characterized by an ENE-WSW to ESE-WNW maximal horizontal stress (oH max = ol or o2)

and a NNW-SSE to NNE-SSW minimal horizontal stress (oH min = o3). This phase therefore

corresponds to an extensional strike-slip tectonic regime. However, in the field, normal faults are clearly

more abundant that strike-slip ones (Figs 6, 11) which would imply a dominantly extensional phase.

144

FRANCOISE BERGERAT ETAL.

Fig. 15.— Zlatnaniva quarry (Berriasian-Barremian limestones).

This site shows NNW-SSE (a) and NNE-SSW (b) compression characterized by strike-slip faulting. The stress tensor

computation for both sets shows a N-S compression (c), but some faults do not Fit with the average stress tensor (due to

the large dispersion in azimuth of each family, especially the left-lateral one: dispersion angle close to 90°).

FlG. 15 .— Analyse de la carriere de Zlatnaniva (calcaires du Berriasien-Barrernien).

Le site montre des compressions NNW-SSE (a) et NNE-SSW (b) caracterisees par des decrochements. Le calcul du

tenseur des contraintes pour la population totale montre une compression N-S (c), mais plusieurs failles tie sont pas en

bon accord avec ce tenseur (en raison de la grande dispersion azimutale de chaque famille, les failles senestres en

particulier montrant un angle de dispersion de pres de 90°).

a b

Fig. 16.— Joint directions in the Miocene (a, 49 measurements) and the Pliocene (b, 34 measurements) in the eastern part of the

Moesian platform

FlG. 16 .— Directions de diaclases dans le Miocene (a : 49 mesures) et le Pliocene (b : 34 mesures) de la partie orientale de la

plate-forme moesienne.

The state of stress characterized by strike-slip faulting and, near the Fore-Balkan, by reverse faulting

with a NNW-SSE to NNE-SSW maximal principal stress ol has been recognized in all formations

ranging in age from the Lower Cretaceous to the Paleocene. At the end of the middle Eocene, the

folding and thrusting of the Stara-Planina and Fore-Balkan area were induced by a major compressional

phase (Ivanov, 1988). The progressive change from North to South (from the Danube river to the front

of the Fore-Balkan) of a strike-slip regime to a purely compressive regime with reverse faulting (Figs 7,

8) suggests that this state of stress characterizes here the same compressional phase.

Does the NNW-SSE to NNE-SSW azimuthal dispersion of the horizontal ol axis (Fig. 9) correspond

to a single roughly N-S compressive state of stress with temporal or local variations or does it

Source: MNHN, Paris

THE MOESIAN PLATFORM: GEODYNAMICAL EVOLUTION

145

Table 5.— Paleomagnetic data from Bulgaria. The columns contain from left to right: Area of sampling; N°, site number on

figure 2; Locality; age of the sampled formation; D(°), mean declination; I (°), mean inclination; authors.

Tableau 5 .— Donnies paleomagnitiques de Bulgarie. De gauche a droite : region d'ichantillonnage, numero du site (N°, voir

localisation figure 2), localite, age de Techantillon, declinaison D(°) et inclinaison I(°), references.

Area

N°

Locality

Age

D(°)

Reference

Forc-Balkan

Smolianovci

Permian-Triassic

18.72

27.15

Nozharov et al., 1980

Bov

Triassic

26.07

39.35

id.

Targoviste

Bajocian-Bathonian

7.4

59.2

Surmont et al., 1991

Stara-Planina

GlNCI 1

Kimmeridgian-Tithonian

7.6

53.8

id.

GlNCl 2

Tithonian-Berriasian

2.9

51.3

id.

Elena

Tithonian-Valanginian

8.6

56.3

id.

Neshkovci

Toarcian

14.5

Kruczyk et al., 1988

KurtHisar

Aalenian

3.7

43.3

id.

Chiflik

Kimmeridgian

30.3

59.5

id.

DolniLom

Bathonian

10.8

57.4

id.

MlTROVCI

Callovian

37.3

59.4

id.

Beli Mel

Callovian

id.

Sredna Gora

Gradec

Bajocian-Callovian

296

Kruczyk et al., 1990

Kraiste

Zablano

Pliensbachian

158

id.

Strandza

Bliznak

Callovian-Kimmeridgian

351

id.

correspond to two distinct stress states? Only one site out of 18 seems to have recorded two distinct

stress states: One NNW-SSE (a 1: az. 341 °), the other NNE-SSW (a 1: az. 210°) (Table 2, Fig. 15).

As previously shown, the state of stress characterized by normal faults and an E-W trending

extension could be attributed to the same tectonic phase, characterized by strike-slip faults and both N-S

trending compression and E-W trending extension. The two states of stress could then be explained by a

simple permutation of stresses ol and g 2 (Fig. 13). However this hypothesis is based solely on

mechanical coherence criteria. It would therefore consist of a tectonic phase characterized by a strike-

slip compressive regime near the chain, changing into a strike-slip extensional regime within the

platform to the North. No chronological criterion invalidates or confirms the synchronism of normal and

strike-slip faults. One notes elsewhere that this extension has practically never been observed west of a

Ruse-Veliko Tarnovo line (Fig. 12), whereas the N-S trending compression was recognized over the

whole platform.

The NNE-SSW trending compression recognized in the Pliocene around Silistra (Fig. 10) would

correspond to a phase syn-or post-Pliocene in age. This phase has been recognized in the Carpathians

(SANDULESCU, 1988) and shown to have a N-S to NNE-SSW horizontal ol (HlPPOLYTE &

SANDULESCU, 1996). The coaxiality of the two phases (Eocene and Pliocene) makes it difficult to know

wheter all the faults measured in the Cretaceous and the Paleocene are related to the Eocene phase ; they

could partly be syn-or post-Pliocene.

Finally, brittle structures other than joints are not present in any of the 8 sites studied in the Miocene

of the eastern part of the platform. Although relatively distributed in direction, these joints nevertheless

146

FRAN^OISE BERGERAT ETAL.

show a well marked E-W maximum (Fig. 16a). One finds this same E-W trend for joints measured in the

Pliocene limestones in Professor Ichirkovo (Fig. 16b), which would imply a quite recent age for these

fractures.

The majority of these joints are extension fractures but small part may be hybrid-fractures. They can

be related to regional stress field and are parallel or subparallel to the direction of greatest horizontal

stress (HANCOCK, 1991). That supposes here a horizontal o3 axis close to N-S. It is difficult to interpret

this recent direction of extension in the framework of the Carpatho-Balkan domain.

CONSTRAINTS FROM PALEOMAGNETISM

The most discussed question regarding the Moesian platform is whether it underwent rotation (and/or

translation). For some authors (e.g. SENGOR et al ., 1984), the Balkan and the Moesian platform have not

undergone any post-Hercynian rotation relative to stable Europe. For others (e.g. HSU et a! ., 1977), they

have undergone an important counterclockwise rotation. The microtectonic analysis alone does not

discriminate between these two hypotheses. Thus we have analyzed the available paleomagnetic data

(Fig. 2). The data considered come from rocks of Permian to Lower Cretaceous age in the Fore-Balkan

(NOZHAROV et al, 1980; SURMONT etal, 1991), the Stara Planina (KRUCZYK et al., 1988; SURMONT et

al ., 1991) and the Sredna Gora, Kraiste and Strandza zones (KRUCZYK et al ., 1990).

The results of analyses undertaken by NOZHAROV et al. (1980) in the Northwestern part of the Fore-

Balkan are compatible with values obtained for the Moesian platform. Similarly, sites analyzed in the

whole Stara Planina by KRUCZYK et al. (1988) show neither relative rotation between its western and

central parts, nor rotation relative to stable Europe.

Paleomagnetic studies realized by SURMONT et al. (1991) in the western part of the Fore-Balkan and

the central and eastern parts of the Stara Planina indicate that these zones have undergone a low

amplitude counterclockwise rotation (about 10°-15°) as compared to the Eurasia. This result was

obtained using the theoretical evolution of the declination of a site centered on the external Balkanides,

assuming it was rigidly linked to Eurasia, and the curve of pole drift of the Eurasia established by

WESTPHAL et al. (1986). The use of the synthetic curve proposed for Eurasia by these authors lead to a

much smaller rotation. According to SURMONT et al. (1991) the rotation occurred during the Eocene

phase.

The investigations led by KRUCZYK et al. (1990) on Jurassic sediments of Sredna Gora, Kraiste and

Strandza, and compared to reference data for the European platform imply a counterclockwise rotation

of southern Bulgaria of approximately 10° to 20°. These results compared to those obtained on

magmatic rocks of the Sredna Gora, of Santonian-Campanian age (NOZHAROV et al ., 1987), suggest that

Jurassic rocks have acquired their post-folding magnetic remanence at the end of the Upper Cretaceous

(KRUCZYK et al., 1990). The difference between the magnetic declination and the geomagnetic field

expected for the Upper Cretaceous in Southern Bulgaria (by using the curve of pole drift of WESTPHAL

et al., 1986) indicates a counterclockwise rotation of 10° to 20° with respect to Eurasia.

If one considers the all of these results together, despite differences of interpretation according to

authors, the whole Moesian platform, Fore-Balkan and Stara Planina area, including even the Sredna

Gora, has undergone a small or null rotation (maximum counterclockwise rotation: 20°) by reference to

stable Europe. This postulated rotation may have occurred at the end of the Upper Cretaceous, as

suggested by KRUCZYK et al. (1990), since the direction of the Eocene compression characterized by the

analysis of microtectonic data in the Moesian platform is close to that recognized, for the same period,

immediately westward in the Pannonian Basin (FODOR et al., 1992), as well as in the whole western

European platform (LETOUZEY et al., 1986; BERGERAT, 1985, 1987).

A major rotation of the Moesian-Balkan domain did not occur according to existing paleomagnetic

data.

A study of Mesozoic series of the Unit of Bihor (Apuseni Mountains, Romania) by SURMONT (1989)

has shown a remagnetization in the Upper Cretaceous-Paleocene. During and after this remagnetization,

the Bihor Unit has undergone a clockwise rotation of approximately 60° as compared to the Eurasia, that

SURMONT (1989) interprets, at least partly, as due to the pushing of Moesia. This Unit belongs to the

South-Pannonian block (Tisza), as well as the Mecsek and Villany Mountains in Hungary. The position

of this block at the beginning of the Tertiary is poorly constrained and controversial (BALLA, 1985;

CSONTOS et al., 1992; CSONTOS, 1995). In the present state of our knowledge, one can neither confirm

Source: MNHN. Paris

THE MOESIAN PLATFORM: GEODYNAMICAL EVOLUTION

147

nor invalidate the hypothesis of a westward translation of Moesia, nor constrain its role in the rotation of

the Tisza.

CONCLUSION

No compressive event recorded in the Moesian platform can, by its direction or by its age,

correspond to the major phase that has affected parts of the Sredna Gora and the Strandza-Eastern

Rhodope during the Albian.

The oldest tectonic phase that seems to have affected the Moesian platform since the beginning of the

Cretaceous is the post-Aptian and pre-Maastrichtian phase with an E-W to ESE-WNW trending

compression and a N-S to NNE-SSW trending extension. Although characterized both by strike-slip and

normal faults, the abundance of normal faults and their systematic presence in most sites show that this

phase is dominantly extensional. It is thus tempting to relate deformation of the Moesian platform to

extension that occurred since the Senonian in the Eastern Stara Planina (flysch of Emine).

The major phase dating from the end of the middle Eocene (Ivanov, 1988), has been recorded in all

rocks of Cretaceous and Paleocene age of the Moesian platform. This phase corresponds to the

formation of significant in the Balkan and the Rhodope. According to NACHEV (1980), the formation of

structures was generated in two episodes, one after the lower Paleocene and one after the middle

Eocene. One can perhaps see in the distribution of al directions (NNW-SSE to NNE-SSW) the trace of

these two episodes.

Finally, the last recognizable tectonic phase, affecting Pliocene limestones in the Northwestern part

of the platform, corresponds probably to a phase well known in the Romanian Carpathians

(SANDULESCU, 1988) and little mentioned in Bulgaria until now (BERGERAT & PlRONKOV, 1994). This

phase, has a direction of compression close to N-S, that is practically identical to the Eocene phase in the

Bulgarian part of the Moesian platform. In the Southern Carpathians in contrast the same direction of

compression is present in the Late-Miocene and Pliocene, but is quite different in the early-middle

Miocene (HlPPOLYTE & SANDULESCU, 1996). This therefore demonstrates a notable geodynamical

change during the Miocene in the Carpathians area.

The Alpine evolution of the Moesian platform can be characterized in terms of stages, from the

beginning of the Upper Cretaceous to the Pliocene, related both to the opening of the Black Sea and to

the convergence of the Eurasian and Arabo-African plates.

The timing and mechanism of the formation of the Black Sea basin have been long debated and

contradictory models for its origin have been proposed (e.g. ADAMIA et al ., 1974; BOCCALETTI et al .,

1974; BRINKMANN, 1972; HSU et al. 1977; FlNETTI, 1995; FlNETTl et al, 1988; LETOUZEY et al., 1977;

MANETT1 et al, 1988; OKAY et al, 1994; ZONENSHAIN & LE PlCHON, 1986). However, the latest

geological and geophysical data indicate that the western Black Sea has an oceanic nature and confirm

its interpretation as a back arc basin, behind the Rhodope-Pontide arc. Subduction initiated in the

Aptian-Cenomanian of the neo-Tethys ocean northward beneath the Eurasian margin (Fig. 17a) caused

the development of the arc and the basin (SENGOR & YlLMAZ, 1981; GORUR, 1988; OKAY et al ., 1994).

Using this model, some authors (HSU et al, 1977; OKAY et al, 1994) proposed that the Senonian Sredna

Gora rifting is related to that of the Black Sea (Fig. 17a).

The extensional phase (a3 NNW-SSE to NNE-SSW) in the Moesian platform, post-Aptian-pre-

Maastrichtian in age, may also be related to this extension. Some of the normal faults produced during

this extensional event, are preserved and evident on seismic reflection lines of the Moesian platform area

where one can see that the normal fault activity terminated before the end of the Cretaceous (FlNETTl et

al, 1988).

At the end of the Paleocene, the opening of the western Black Sea was completed.

The consumption of the Tethyan oceanic crust produced an important compressive tectonic event that

affected the Great Caucasus, Crimea and Balkanides (FlNETTl, 1995).

The Rhodope fragment, collided with the Southern edge of the Eurasian plate inducing the

development of nappes originating from the Rhodope basement, followed by thrusting farther to the

North in the Balkan (Fig. 17b).

This compressive phase, of middle Eocene age in the Balkan, has also been recorded in the Moesian

platform, dominantly by strike-slip faulting in a NNW-SSE to NNE-SSW compressive stress field.

148

FRANgOISE BERGERAT ETAL.

SP - FB - M

Fig. 17.— Upper Cretaceous (A) and middle Eocene (B) geodynamical stages of the Balkanides (modified after Dercourt in

Durand-Delga etal ., 1988: Ivanov, 1988).

Abbreviations are as follows: R, Rhodope; SG, Sredna Gora; LK, Luda Kamcija; SP, Stara Planina; FB; Fore-Balkan;

M, Moesian platform.

Fig. 17 .— Giodynamique des Balkanides au Cretace superieur (A) et a I'Eocene moyen (B) (modifie d'apres Dercourt in

Durand - Delga et al., 1988; Ivanov, 1988).

R : Rhodope, SG : Sredna Gora, LK : Luda Kamcija, SP : Stara Planina, FB : Prebalkan, M : plate-forme moesienne.

Finally, continuing convergence caused compression, as shown by the small reverse and strike-slip

faults yielding a NNE-SSW compression in the Pliocene of the Moesian platform.

ACKNOWLEDGEMENTS

This study was carried out as part of the Peri-Tethys program. Special thanks are given to Z. IVANOV

for his constructive observations and to P. HANCOCK for his review. We acknowledge V. CARRERE and

M. GORDON for improving the English of this paper.

REFERENCES

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problem of formation of deep basin of the Black Sea. Geotectonika, 1: 78-94 (in Russian).

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THE MOESIAN PLATFORM: GEODYNAMICAL EVOLUTION

149

ANGELIER, J., 1990.— Inversion of field data in fault tectonics to obtain the regional stress. II: a new rapid direct inversion

method by analytical means. Geophysical Journal International, 103: 363-376.

Balla, Z., 1984.— The Carpathian loop and Pannonian basin: a kinematic analysis. Geophysical Transactions, 30 (4): 313-

353.

BERGERAT, F., 1985.— Deformations cassantes et champs de contrainte tertiaires dans la plate-forme europeenne. These de

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Source: MNHN, Pans

Scythian Platform: chronostratigraphy

and polyphase stages of tectonic history

Anatoly M. NlKISHlN"', Sierd CLOETINGH Sergey N. BOLOTOV ",

Evgenij Yu. BARABOSHKIN"', Ludmila F. KOPAEVICH",

Bronislav P. NAZAREVICH Dmitry I. PANOV 1 ' 1 ,

Marie-Frangoise BRUNET 131 , Andrei V. ERSHOV" 1 , Vera V. I LINA

Svetlana S. KOSOVA 141 & Randell A. STEPHENSON 121

0> Geological Faculty, Moscow State University, 119899 Moscow, Russia

<2 ' Vrije Universiteit, De Boelelaan 1085, 1081 HV Amsterdam, The Netherlands

(,) Departement de Geotectonique, CNRS URA 1759, case 129, Universite P. et M. Curie

75252 Paris Cedex 05, France

<4 ' Central Geophysical Expedition, Moscow. Russia

ABSTRACT

The tectonic history of the Scythian Platform, overlying the pre-Mesozoic folded basement, is presented on the basis of

chronostratigraphic and subsidence history data. The following stages can be recognized: Early Triassic-Camian - origin of

continental rift system; pre-Norian-Hettangian - syncollisional orogeny; Sinemurian-early Aalenian - synrift uplift; late

Aalenian-pre-Callovian - syncompressional regional subsidence with pre-Callovian uplift; Callovian-Tithonian - synrift

subsidence and uplift; early Berriasian - congressional deformations; late Beiriasian-Barremian - regional subsidence; Aptian-

Albian - continental rift system origination; Late Cretaceous-Eocene - thermal regional subsidence complicated by stress

events; Oligocene-early Miocene - syncompressional regional subsidence; middle Miocene-Quaternary - syncollisional molasse

basin formation with five main compressional phases.

RESUME

La plate-forme Scythe : chronostratigraphie et histoire tectonique polyphasee.

L'histoire tectonique de la plate-forme Scythe, reposant sur un socle plisse ante-mesozoique est presentee sur la base de

donnees de reconstitution de la chronostratigraphie et de la subsidence. On peut reconnaitre les etapes suivantes : Trias

inferieur-Carnien - origine du systeme de rift continental; pre-Norien-Hettangien - orogenese syncollisionelle ; Sinemunen-

Aalenien inferieur - soulevement synrift; Aalenien superieur-pre-Callovien - subsidence regionale syncompression

accompagnee d’un soulevement d’age pre-Callovien; Callovien-Tithonique - subsidence synrift et soulevement; Berriasien

inferieur - deformations en compression ; Berriasien superieur-Barremien - subsidence regionale ; Aptien-Albien - origine du

systeme de rift continental ; Cretace superieur-Eocene - subsidence regionale d’origine thermique, compliquee par des

variations des contraintes tectoniques ; Oligocene-Miocene inferieur - subsidence regionale syncompression ; Miocene moyen-

Quaternaire - formation de bassins molassiques syncollisionels avec cinq phases principales de collision.

NlKISHlN, A.M., CLOETINGH, S., Bolotov, S.N., Baraboshkin, E.Y., Kopaevich, L.F., Nazarevich, B.P., Panov, D.I.,

Brunet, M.-F., Ershov, A.V., Ilina, V.V.. Kosova, S.S. & Stephenson, R.A., 1998.— Scythian Platform: chrono¬

stratigraphy and polyphase stages of tectonic history. In: S. Crasquin-Soleau & E. Barrier (eds), Pcri-Tethys Memoir 3:

stratigraphy and evolution of Peri-Tethyan platforms. Mem. Mus. natn. Hist, nat ., 177 : 151-162. Paris ISBN : 2-85653-512-7.

Source: MNHN. Paris

152

ANATOLY M. NIKISHIN ETAL.

Peri-Caspion

- basin-

EAST-EUROPEAN

PLATFORM

MOESIAN

PLATFORM

ARABIAN

PLATFORM

Fig. 1.— Tectonic position of the Scythian Platform. 1-3. Great Caucasus-Southern Crimea-Dobrogea orogen: 1, pre-Mesozoic

basement; 2. intensively folded / thrusted Mesozoic and (or) Cenozoic; 3. gently deformed Mesozoic and (or) Cenozoic

(deformed margins of the Scythian Platform); 4, Late Cenozoic molasse basin; 5, basin with oceanic or very thinned

continental crust; 6. thrust zone; 7, zone of continental subduction in Zagros. Some basins and other tectonic units: Az -

Azov Basin (recent Azov Swell), Da - Dagestan basin. Do - Dobrogea orogen, EM - East Manych basin, EPC - Eastern

Pre-Caucasus basin (includes western part of the Terek-Kuban basin, Mozdok basin and in pre-Neogene times -

Dagestan basin), IKB - Indol-Kuban molasse basin, K - Kayasula basin, KB - Karkinitsky (Northern Crimea) basin, KS

- Karpinsky Swell, M - Mozdok basin, NF - Novo-Fedorovsk basin, SH - Stavropol High, TCB - Terek-Caspian

molasse basin, WPC - Western Pre-Caucasus basin (includes Indol-Kuban molasse basin).

FlG. 1 .— Carte tectonique de la Plate-forme Scythe. 1-3, Orogene Grand Caucase-Crimee meridionale-Dobrogee : 1, socle

ante-Mesozotque ; 2, Mesozoique et (ou) Cenozoique intensement plisse / chevauche ; 3, Mesozoique et (ou)

Cenozoique faiblement plisse / chevauche (marges deformees de la Plate-forme Scythe); 4, bassin molassique

Cenozoique superieur; 5. bassin avec une croute oceanique ou une croute continentale tres amincie ; 6, zone de

chevauchement; 7. zone de subduction continentale dans le Zagros. Quelques bassins et autres unites tectoniques : Az -

bassin d'Azov (recemment ride d'Azov), Da - bassin du Dagestan, Do - Dobrogee, orogene, EM - bassin du Manych

oriental, EPC - bassin oriental du Pre-Caucase (comprenant la partie occidentale du bassin Terek-Kuban, le bassin de

Mozdok et a lante-Neogene - le bassin du Dagestan), IKB -bassin molassique d'Indol-Kuban, K - bassin de Kayasula,

KB - bassin de Karkinitsky (Crimee septentrionale), KS - ride de Karpinsky, M - bassin de Mozdok, NF - bassin de

Novo-Fedorovsk, SH - Haut de Stavropol, TCB - bassin molassique Terek-Caspien, WPC - bassin occidental du Pre-

Caucase (comprenant le bassin molassique dlndol-Kuban).

Source: MNHN. Paris

SCYTHIAN PLATFORM: CHRONOSTRATIGRAPHY AND TECTONIC HISTORY

153

INTRODUCTION

The Scythian Platform is located between the Great Caucasus-South Crimea orogen to the South, and

the East European Platform to the North (Fig. 1). It includes the Scythian Platform itself underlain by a

Late Paleozoic folded basement and covered by the Terek-Caspian and Indol-Kuban foreland molasse

basins of the Great Caucasus-Southern Crimea orogen.

This region is relatively well investigated (MlLANOVSKY & Khain, 1963; MlLANOVSKY, 1968;

CHEKUNOV et al ., 1976; KORONOVSKY et at ., 1987; GOFMAN et al. y 1988; ROSTOVTSEV, 1992;

Letavin, 1980; KUNIN et al ., 1989). The thickness of sedimentary basin cover is typically more than 5

km and up to 12 km. The following lithostratigraphic complexes make up the sections of the Scythian

Platform: Triassic, Jurassic-Eocene, Oligocene-lower Miocene, middle Miocene-Quaternary.

CHRONOSTRATIGRAPHY OF THE SCYTHIAN PLATFORM

The stratigraphy of the Scythian Platform is based on the data from a few hundreds of wells.

These data are summarized on the chronostratigraphic profiles located in the eastern Pre-Caucasus

area and Crimea area (Figs 2-4).

Early-Middle Triassic-early Carnian: Lower-Middle Triassic and lower Camian strata are located in

a few rift-type basins and were folded and partially eroded in pre-Norian and pre-Aalenian-Bajocian

times. Two main sedimentary basins, the East-Pre-Caucasus basin and West-Pre-Caucasus-Crimea basin

Stavropol

Azov Sea

Krasnodar

Groznyy

Vladika

Fig. 2.— Location map. Solid lines: Cau. - location of the chronostratigraphic profile for the eastern Pre-Caucasus region (see

figure 3), Cri. - location of the chronostratigraphic profile for the Central-Northern Crimea (see figure 4). Numbers -

location of wells used for backstripping analysis (see figure 5).

FlG. 2 .— Carte de localisation. Trait plein: Cau. - localisation du profit chronostratigraphique de la region orientate du Pre-

Caucase (voir figure 3), Cri. - localisation du profit chronostratigraphique de la region cent rale du nord de la Critnee

(voir figure 4). Chiffres - localisation des forages utilises pour Vanalyse du “ backstripping " (voir figure 5).

154

ANATOLY M. NIKISHIN ET AL.

were separated by the Central-Pre-Caucasus (Stavropol) High. The East-Pre-Caucasus basin was

represented by a water covered rift system. It is possible to recognize at least three paleorift troughs of

East-West trend: East Manych to the North, Mozdok to the South and Kayasula in the middle

(Nazarevich et al ., 1986; NIKISHIN et al ., 1994). A few remnants of other rift valleys can be proposed

too. The East Manych trough was a shallow and then a relative deep-water trough with clay, marl,

carbonate uncompensated sedimentations and volcanic centres. Kayasula trough had a same

paleogeography, but was shallower and carbonate sedimentation dominated together with volcanism.

Underwater high between East Manych and Kayasula troughs was covered by carbonatic reef belt.

Preliminary data of flysch-like sediments in the Mozdok basin suggest that this trough was a part of an

Early-Middle Triassic passive margin.

Reliable stratigraphic data on the West-Pre-Caucasus-Crimean area are absent. It was a marine

sedimentary basin with Northern-Crimea-Azov and Novo-Fedorovsk rift-like troughs with carbonate,

clay and clastic depositions (Slavin, 1986; BOIKO, 1993; BOIKO et al ., 1986). The Lower-Middle

Triassic of the West-Kuban (West-Pre-Caucasus) basin is represented by carbonates and clastic deposits.

The paleowater-depth increased to the direction of recent Great Caucasus (BOIKO et al., 1986). Possibly

uplift, erosion and deformation took place in a pre-Norian time.

Late Triassic: three main Late Triassic sedimentary basins in the Scythian Platform can be

recognized: the East-Pre-Caucasus or Nogaisk basin, West-Pre-Caucasus or West Kuban basin and the

Crimea region. The Nogaisk basin originated at the site of the former Early-Middle Triassic rift system.

It covers East Manych-Kayasula-Mozdok rifted basins. The Nogaisk basin was a continental

volcanogenic-molasse basin of the Norian-Rhaetian age (NAZAREVICH et al., 1986). It was filled by

sands, silts, conglomerates and volcanics (andesites, rhyolites, ignimbrites, tuffs). The Karpinsky Swell

zone was upthrusted to the Nogaisk basin (NAZAREVICH et al., 1986; NIKISHIN et al., 1994; SOBORNOV,

1995). The Great Caucasus region and Karpinsky Swell zone were main sources of clastic material. The

Central Pre-Caucasus (Stavropol) area was an uplifted (mountain-?) region. The West-Pre-Caucasus

basin underwent structural changes; flysch, bioherm and shallow-water clastic sediments formed in

different subzones of the former basin (BOIKO, 1993; PRUTSKY& LAVRISCHEV, 1989). At the margin of

the Great Caucasus basin a carbonate reef belt originated (BoiKOef al., 1986). The former Azov-

Northern-Crimea rift basin suffered inversion tectonics. The data on Central-Northern Crimea are very

poor (Slavin, 1986); in a few wells Late Triassic dacites, andesites and intrusions of diorites have been

drilled.

In all these basins Late Triassic complexes unconformably overlie older sediments. The Late Triassic

is mainly represented by Norian-Rhaetian. The presence of the early Carnian is more local, restricted

probably to Carnian marine sediments. Hettangian sediments are not documented in the Scythian

Platform. It appears that the main orogenic deformations took place in the late Carnian-pre-Norian time,

at the Rhaetian / Hettangian boundary and in Hettangian.

Early Jurassic: the Scythian Platform was an uplifted area during the Hettangian, probably as a result

of syncompressional tectonics. The Sinemurian is absent as well, possibly due to thermal synrift uplift

(the rifting took place in the zone of recent Great Caucasus). Sinemurian sedimentation took place along

the Great Caucasus trough. Pliensbachian and Toarcian strata are located in a few rift-like troughs with a

typical width of about 30 km and thickness of sediments (mainly claystones and siltstones with possible

Pliensbachian dacites and rhyolites) about 150-200 meters (STAFEEV et al., 1995).

Aalenian-Bathonian: during the late Aalenian-early Bathonian a regional subsidence of the Scythian

Platform took place (KORONOVSKY et al., 1987; Panov & GUSCHIN, 1987). The main sedimentary

basins were the following: the Kalmyksky basin (in the Karpinsky Swell zone), East-Pre-Caucasus

Fig. 3.— Chronostratigraphic profile for the Eastern Pre-Caucasus region, for location see figure 2 (solid line Cau.). Lithology

is simplified: 1. limestones; 2, marls mainly; 3, clays; 4. siltstones; 5, sandstones; 6, conglomerates; 7, evaporites; 8,

deep-water turbidites; 9, basalts mainly; 10. volcanic arc complex (volcanites, flysch, clastic sediments); 11, acidic

volcanites and molasse; 12, erosional boundary; 13, orogenic event with folding / thrusting and uplift; 14, rift phase; 15,

compressional phase.

FlG. 3. — Profit chronostratigraphique de la region orientate du Pre-Caucase, position figure 2 (ligne Cau.). La lithologie est

simplifiee : I. calcaires; 2, marnes ; 3. argiles ; 4, microgres ; 5, gres ; 6, congtomerats ; 7. evaporites ; 8. turbidites

d'eau profonde ; 9. principalement basaltes ; 10, complexe d'arc volcanique (volcanites, flysch, sediments clastiques);

11. volcanites acides et molasse ; 12, limite d'erosion ; 13, evenement orogenique avec plissement / chevauchement et

soulevement; 14. phase extensive ; 15, phase compressive.

Source: MNHN. Paris

SCYTHIAN PLATFORM: CHRONOSTRATIGRAPHY AND TECTONIC HISTORY

155

Source: MNHN. Paris

156

ANATOLY M. NIKISHIN ETAL.

Fig. 4.— Chronostratigraphical profile for the Central-Northern Crimea, for location see figure 2 (solid line Cri.). Lithology is

simplified. For legend, see figure 3.

FlG. 4 — Profit chronostratigraphique de la region centrale du nord de la Crimee, position figure 2 (ligne Cri.); la lithologie

est simplifiee. Pour la legende voir figure 3.

Source: MNHN, Pans

SCYTHIAN PLATFORM: CHRONOSTRATIGRAPHY AND TECTONIC HISTORY

157

basin, East-Kuban basin and West-Kuban basin in the West-Pre-Caucasus area, and the Southern-

Crimea-Great Caucasus deep-water trough (orogenic zone). All the basins were Tilled by clay and silt

deposits mainly. During late Aalenian-Bajocian-earliest Bathonian time, main epoch of the Scythian

Platform subsidence occurred. It was a time of major subduction-related magmatism in the

Transcaucasus area and orogenic activity inside the Great Caucasus belt. The East-European Craton was

the main source of the clastic material supply. The basins water-depths were not deep, and coal-bearing

sediments originated in several places. The late Bathonian-early Callovian is mainly a gap in the

stratigraphy being a time of syncompressional uplift and minor erosion.

Callovian-Tithonian: the Callovian-Late Jurassic was the next epoch of the Scythian Platform

history. Although sedimentary basins were nearly the same as in the Middle Jurassic, configurations and

structure of the basins changed. Shallow-water sedimentation dominated during the Callovian-early

Oxfordian, with deposition of clays, silts, conglomerates and carbonates. A few more deep areas were

located in the central parts of the basins. In the middle-late Oxfordian, the paleogeography was nearly

the same. Reef belts began to grow around the deep parts of the basins, and uncompensated

sedimentation took place in the central portions of the basins. In the Kimmeridgian-Tithonian, after the

changes of basin configuration and uplift of the southern marginal zone of the Scythian Platform,

evaporite sedimentation dominated. This phase marks the most important evaporite event in the Scythian

Platform history. Callovian-Upper Jurassic sedimentations had a great diversity and were very

heterogeneous. At the end of the Tithonian and in the Berriasian, a change of paleogeography took place

together with a weak uplift and compressional tectonics (MlLEEV et al. , 1995). An angular unconformity

can be recognized in a few seismic profiles at the Jurassic / Cretaceous boundary or intra-Berriasian.

Early Cretaceous: the Early Cretaceous was the following stage of the platform cover formation. It

was a time of unstable tectonic environments inside the Scythian Platform. The following sedimentary

basins can be recognized: the East-Pre-Caucasus (Dagestan) basin, West-Pre-Caucasus (Kuban) basin,

Northern Crimea (Karkinitsky) basin, South-Western Crimea (Western Pre-Yaila) basin, South-Eastern

Crimea (Eastern Pre-Yaila) basin, the Salgir Graben in the centre of-the Southern Crimea, and the Great

Caucasus deep-water trough. Carbonate and terrigenous shallow-water sediments dominated in the Early

Cretaceous of the Scythian Platform. We can recognize two main substages of history in the Early

Cretaceous: late Berriasian-Barremian, and Aptian-Albian* During the first substage shallow-water

sedimentation dominated but although the tectonic setting is not clear, a tensional regime is probable.

The Aptian-Albian was a time of continental rifting, but more accurate stratigraphicai and structural data

are required to recognize different events in this history. The Salgir Graben in the Southern Crimea

originated in the Aptian-early Albian (MURATOV, 1969). Karkinitsky (Northern Crimea) deep graben

(more than 3 km of sediments) was active in Aptian-Albian with volcanism in the middle-late Albian

(Muratov, 1969; CHEKUNOV et al ., 1976; Grigorieva et al ., 1981), but subsidence started in the

Barremian or late Hauterivian. Belogorsk (or Eastern-Pre-Yaila) graben (more than 1 km of sediments)

originated in the Aptian-Albian in the south-eastern portion of the Crimea. Kuban basin and Dagestan

basin suffered rapid subsidence. Volcanism took place in a number of regions in the middle-late Albian-

earliest Cenomanian (Grigorieva et al. , 1981).

Late Cretaceous: the Late Cretaceous consists of a carbonate platform cover on the Scythian

Platform and the southern part of the Russian Platform. The relatively deepest parts of the basin were

close to the Great Caucasus trough and overlying some Early Cretaceous rifted basins (for example - the

Karkinitsky basin). The Great Caucasus trough was the deep-water basin with a flysch sedimentation

mainly.

Paleocene-Eocene: the Paleocene-Eocene paleogeography was nearly the same as in the Late

Cretaceous. Carbonate shallow-water sedimentation dominated round the Scythian Platform. The

deepest portions of the basins were close to the Great Caucasus trough, and some changes of basins

configuration can be observed in the Platform, pointing to some changes of tectonics in the Paleocene

and at the Paleocene / Eocene boundary.

Oligocene-early Miocene (Maykop time): during the Oligocene - early early Miocene time, the

Maykop formation had been formed. The Maykop formation is widespread in areas of the Scythian

Platform, Great Caucasus trough, Black Sea basin, and the Turan Platform. It has been formed in the

large marine basin known as Paratethys (MlLANOVSKY, 1968; SCHERBA, 1987; KORONOVSKY et al .,

1987; POPOV et al ., 1993). The early Oligocene basins sometimes cross outlines the former basins and

follow closely to the configurations of younger molasse basins.

158

ANATOLY M. NIKISH1N ET AL.

The thickness of Maykop formation is up to 1-3 km in the Scythian Platform, filling a deep-water

basin with water-depths up to 400-1000 meters. The Maykop formation is represented mainly by clay

and sandstones, including series of clinoforms southwardly dipping toward the Caucasus (KUNIN et al.,

1989). The water depths have increased to the south. Deep water basin originated very rapidly in the

early Oligocene. Limited evidence exists that at the northern slope of the Great Caucasus clinoforms dip

to the north, and there are a few proposed latest Eocene-early Oligocene thrusts to the Scythian Platform

from the Caucasus belt.

Extensional faults at the bottom of Maykop formation cannot be recognized on the seismic profiles

(KUNIN el al., 1989). Deep-water basins have formed mainly due to vertical subsidence without

significant role of faults. Gravitational olistostrome sedimentary bodies moving to the inner portions of

the Maykop basin are well established on seismic profiles and field outcrops in the lower part of

Maykop formation (SCHERBA. 1987; KUNIN et al., 1989).

Late early Miocene-Quaternary: in the late early Miocene-early middle Miocene (Tarkhanian-

Konka, 17-13.7 Ma ago) the Scythian Platform was filled by marine and continental sediments

(MlLANOVSKY, 1968). Uplift evidence exists of a central part of the Great Caucasus since the

Chokrakian (16.5 Ma ago) (GONTSHAROVA & SCHERBA, 1996). In the Sarmatian time (late middle

Miocene - early late Miocene) a large supply of molasse sediments started from the Great Caucasus. In

the late Sarmatian time, the Great Caucasus existed as mountains (MlLANOVSKY, 1968; KUNIN et al.,

1989). The uplift of Great Caucasus was irregular in time during the Miocene, Pliocene and Quaternary;

phases of rapid uplift alternated with tectonically more quiet epochs (MlLANOVSKY, 1968).

Zones with a different geological history are seen in the eastern part of North Caucasus basin in

Miocene-Quaternary. Examples are the Terek-Caspian basin, East-Manych basin, zone of Karpinsky

Swell. Stavropol High. The Terek-Caspian basin underwent subsidence up to 2-4 km and was filled by

molasse sediments. These molasses make up large scale clinoforms of alluvial to marine sediments.

The Karpinsky Swell zone has undergone an uplift up to 0.5 - 1 km in Miocene; at the end of

Pliocene - Quaternary time its eastern part has subsided and was covered by a sedimentary cover, at the

same time its western part has continued slow uplift (MlLANOVSKY, 1968).

The East Manych basin zone was a topographically low depression filled by molasse in Pliocene-

Quaternary time. The molasse clinoforms progradated toward this basin both from the Great Caucasus

Mountains to the south and Karpinsky Swell to the north. The Stavropol High has undergone slow uplift

in Miocene-Quaternary.

ONE DIMENSION BURIAL HISTORY

MODELLING OF THE SCYTHIAN PLATFORM

We used computer modelling to investigate the burial history of the Scythian Platform (NlKISHIN et

al., 1994) using the data from 128 wells and sections. A number of curves of tectonic subsidence rates

are shown on figure 5. The analysis of the model shows the following features:

Early-Middle Triassic-Camian: rapid subsidence event for many wells; possible interpretation: rift

phase and post-rift subsidence;

Triassic / Jurassic boundary, uplift event for a few wells; possible interpretation - inversion

compressional tectonics;

End of the Early Jurassic : a few wells show rapid subsidence event at this time; possible

interpretation - weak rift phase coeval with main phase of opening of the Great Caucasus trough;

Late Aalenian-Bajocian-early Bathonian: subsidence phase; possible interpretations:

syncompressional foreland-type subsidence or postrift subsidence accelerated by compression;

Pre-Callovian changes of subsidence curves: possible compressional event;

Late Jurassic: rapid subsidence, possible extensional phase;

Jurassic / Cretaceous boundary: changes of subsidence curves; possible compressional event;

Cretaceous-Paleocene-Eocene: gentle thermal subsidence accelerated in Mid-Cretaceous time in

many areas by possible tension events; changes of subsidence curves at the Cretaceous / Paleocene and

Paleocene / Eocene boundaries for many areas show proposed changes in a stress field;

Source: MNHN, Paris

SCYTHIAN PLATFORM: CHRONOSTRATIGRAPHY AND TECTONIC HISTORY

159

Oligocene : rapid subsidence phase, syncompressional subsidence;

Middle Miocene-Quaternary : five rapid subsidence events: 16.6-15 Ma ago (Tschokrakian-

Karaganian); 12.4-9.7 Ma ago (Sarmatian); 7-5 Ma ago (at the Meotian / Pontian boundary-Pontian);

3.7-1.8 Ma ago (Akchagylian); 1.6-0 Ma; these rapid subsidence phases coincide with the main

deformational compressional phases in the Great Caucasus; hence it appears that they are related to

collisional phases; some areas underwent syncompressional uplift during the late phases of collision.

TECTONIC HISTORY OF THE SCYTHIAN PLATFORM

The proposed tectonic history of the Scythian Platform is based on geological data (including data

from the Great Caucasus-Southern Crimea orogen and other surrounding areas) and inferences from the

computer burial history. The short scenario of the proposed history is represented below (Figs 3-4).

Early-Middle Triassic-Carnian: a major continental rift system covered by post-rift cover originated

along the former Late Paleozoic orogen; it includes the East Manych basin, Kayasula basin, Mozdok

basin, Kiliya-Karkinitsky-Azov belt of basins, Tulcea (Dobrogea) basin, rift basins of the Moesian

platform and Turan platform, activated Dnieper basin. Rift related magmatism took place in many rifted

basins. Compressional tectonics took place in pre-Norian time.

Norian-Rhaetian. Compressional tectonics, origin of molasse basins, orogenic magmatism.

Triassic / Jurassic boundary and Hettangian: collisional orogeny (possible collision with

Transcaucasus terrane, Iran terrane, Western and Eastern Pontides terranes).

Sinemurian-early Aalenian: high stand of the Scythian Platform, origin of a Great Caucasus deep¬

water rifted trough and weak rifting in the uplifted Scythian Platform.

Late Aalenian-early Bajocian: collisional orogenic events in the Crimea, weak compressional events

in the Great Caucasus and regional subsidence of the Scythiaji Platform.

Late Bajocian-Bathonian: during the Bajocian a large subduction-related magmatic belt was active in

the Transcaucasus-Pontides area. Intra-Bajocian inversion tectonics took place along the Great

Caucasus-South Crimean belt. This inversion tectonics continued to Bathonian and had a culmination at

the end of Bathonian-beginning of the Callovian as an orogeny along the Great Caucasus-South Crimean

belt. Regional subsidence of the Scythian Platform took place in Bajocian-early Bathonian and was

terminated by pre-Callovian uplift.

Callovian-mid-Berriasian: a new rift phase took place in the Callovian-Tithonian along the Great

Caucasus-South Crimean belt, in the Dobrogea and the Moesian Platform. The rifting was accompanied

by subsidence of the Scythian Platform.

Berriasian: orogenic events, uplift and thrusting-folding took place in the Crimea and Scythian

Platform.

Late Berriasian-Barremian: a new system of sedimentary basin configuration originated in the

Scythian Platform, possibly in response to a new weak tension phase. Uplift in pre-Aptian time.

Aptian-Albian: a system of rifted basins originated in the Scythian-Black Sea-Caucasus area.

Late Cretaceous: postrift subsidence and stress events in the Scythian Platform and Great Caucasus.

Paleocene-Eocene: regional thermal subsidence of the Scythian Platform modulated by stress events.

Eocene / Oligocene boundary: rapid syncompressional subsidence of the Scythian platformand

orogenic activity in the Caucasus region.

Oligocene-early Miocene (Maykop time): rapid syncompressional subsidence of the southern part of

the Scythian Platform led to origin of deep-water basins which were filled by clay and sandstones. At

the end of the Maykop time relative compensation by sediments took place in the Scythian Platform.

Middle Miocene-Quaternary: Arabia-Europe continent-continent collision led to collision tectonics

and orogeny in the Caucasus-South Crimea region, and a few molasse basins originated to the north and

south of the Great Caucasus-South-Crimean orogen. The North Caucasus Indol-Kuban and Terek-

160

ANATOLY M. NIKISHIN ETAL.

Geological Time (Ma)

-245 -225 -205-185 -1,65-1,«5-125-105 -85 -65 -45 -25 -5

pg irprq

Ceological Time (Ma)

-245 -325 -305 -185 -165 -145 -125 -105 -85 -65 r 45 -}?5 -5

Indol'skaya-3 well (l)

-r'd

FS,<h

■ST

:j° i .

:-ioc

-20

( 125 )

o!L

J Kayasulinskaya-6 well (6)

20 3

i° N

0 ?

-io<-

(202)

Georgievskaya well (16)

Fig. 5.— Comparison of the tectonic subsidence rate curves for some deep wells on the Scythian Platform. Location of wells

(numbers) are given in figure 2.

FlG. 5 — Comparison des taux de subsidence tectonique pour quelques forages profonds de la Plate-Forme Scythe. La

position des forages (chiffres) est donnee figure 2.

Source: MNHN. Pans

SCYTHIAN PLATFORM: CHRONOSTRATIGRAPHY AND TECTONIC HISTORY

161

Caspian molasse basins underwent five rapid subsidence phases: 16.5-15, 12.4-9.7, 7-5, 3.7-1.8 and 1.6-

0 Ma ago. These times coincide with the main phases of thrusting in the Caucasus. Since the middle

Miocene a syncompressional bulge has been uplifted along the belt of Karpinsky Swell-Donets Basin-

Ukrainian Shield.

CONCLUSIONS

The Mesozoic-Cenozoic subsidence history of the Scythian platform was very complex. It was

controlled by many synrift subsidence or uplift phases, postrift thermal subsidence epochs,

syncompressional phases and epochs of subsidence or uplift (see figures 3, 4). The recent platform cover

is an integrated result of a succession of different tectonic processes with a complex thermal history.

ACKNOWLEDGEMENTS

The work was funded by the Peri-Tethys program. International programmes IGCP-369, INTAS,

EUROPROBE and LITHOSPHERE supported our communications and discussions. Russian Geological

Survey (ROSKOMNEDRA) sponsored our many-years field works. We thank E.E. MlLANOVSKY, V.E.

Khain, A.S. Alekseev, J. Dercourt, P. Ziegler, J.-P. Cadet, V.G. Kazmin, W. Cavazza, B.

FOUKE. N.V. KORONOVSKY, A.N. STAFEEV and E. GRADINARU for fruitful discussions. Netherlands

Research School of Sedimentary Geology contribution n° 960412.

REFERENCES

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Gas Potential of the Cretaceous Sediments of the Southern Ukraine. Naukova Dumka, Kiev: 1-140 (in Russian).

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Mineral Resources of the Great Caucasus. Nauka, Moscow: 147-174 (in Russian).

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Seismostratigraphic Analysis (on the Example of the Eastern Peri-Caucasus Region). VNIIOENG, Moscow: 1-43 (in

Russian).

Letavin, A.I., 1980.— Basement of the Young Platform of the Southern USSR. Nauka. Moscow: 1-150 (in Russian).

Milanovsky, E.E., 1968.— Neotectonics of the Caucasus. Nedra, Moscow: 1-638 (in Russian).

Milanovsky, E.E. & Khain, V.E., 1963.— Geological Structure of the Caucasus. Moscow University Press. Moscow: 1-357

(in Russian).

Mileev, V.S., Rozanov, S.B.. Baraboshkin, E.Yu.. Nikitin, M.Yu. & Shalimov. I.V.. 1995.— The position of the Upper

Jurassic deposits in the structure of the Mountain Crimea. Bulletin MOIP. Geologia. 70 (1): 22-31 (in Russian).

Muratov, M.V. (ed.), 1969.— Geology of the USSR. VIII: Crimea. (1), Geology. Nedra, Moscow: 1-576 (in Russian).

Nazarevich, B.P., Nazarevich, I.A. & Shvydko. N.I., 1986.— Nogaysk (Upper Triassic) volcanogenic-sedimentary

formation of the Eastern Peri-Caucasus region: composition, structure, and relations with pre- and post-Nogaysk

volcanites. In: Formations of Sedimentary' Basins. Nauka, Moscow: 67-86 (in Russian).

Nikishin, A.M., Bolotov, S.N., Ershov, A.V., Nazarevich, B.P., Cloetingh, S.,Lobkovsky. L.I. & Stephenson, R.,

1994.— Geodynamic analysis of the North-Caucasus sedimentary basin according to the data of burial history computer

modelling. In: Yu.G. Leonov, A.F. Morozov & L.N. Solodilov (eds), Tectonics and Magmatism of the East -

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Panov, D.I. & GUSCHIN. A.I.. 1987.— Structural-formational tectonic subdivisions of the Great Caucasus region for the Early

and Middle Jurassic time, and the local Early-Middle Jurassic stratigraphy. In: E.E. Milanovsky & N.V. Koronovsky

(eds). Geology and Mineral Resources of the Great Caucasus. Nauka, Moscow: 124-139 (in Russian).

Popov, S.V, Akhmetiev, M.A., Zaporozhets, N.I., Voronina, A.A. & Stolyarov, A.S., 1993.— History of the Eastern

Paratethys in the Late Eocene-Early Miocene. Stratigrafiya. Geologicheskaya Korrelyatsiya, 1 (6): 10-39 (in Russian).

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Scherba, I.G., 1987.— Olistostromes and problems of Cenozoic tectonics of the Caucasus. In: E.E. Milanovsky & N.V.

Koronovsky (eds). Geology and Mineral Resources of the Great Caucasus. Nauka, Moscow: 191-200 (in Russian).

Slavin, V.I., 1986.— Geological history of the Crimea peninsula in the Triassic. Bulletin MOIP, Geologia, 61 (6): 46-50 (in

Russian).

Sobornov, K.. 1995.— Structural evolution of the Karpinsky swell, Russia. Comptes Rendus de I'Academie des Sciences,

Paris, 321 (II a): 161-169.

Stafeev, A.N.. Panov, D.I., Yutsis, V.V., Smirnova, S.B. &Gushchin, A.I., 1995.— Eastern Peri-Caucasus Lias: unsolved

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Source: MNHN. Paris

Scythian Platform, Caucasus and Black Sea region:

Mesozoic-Cenozoic tectonic history and dymanics

Anatoly M. NlKISHIN Sierd CLOETINGH

Marie-Frangoise BRUNET 111 , Randell A. STEPHENSON 121 ,

Serguey N. BOLOTOV'" & Andrei V. ERSHOV<"

11 ’ Geological Faculty, Moscow State University, 119899 Moscow, Russia

<2 ‘ Vrije Universiteit, De Boelelaan 1085, 1081 HV Amsterdam. The Netherlands

,3> Departement de Geotectonique, CNRS URA 1759, case 129, Universite P. et M. Curie

75252 Paris Cedex 05, France

ABSTRACT _

Collision orogeny in the Caucasus-Scythian area took place at the Triassic / Jurassic boundary. During the Jurassic-Eocene

times the region was in back-arc tectonic environments with many phases of back-arc extension and compression. At the

Eocene / Oligocene boundary the subduction system was transformed into a zone of collision, and during the Oligocene-

Quaternary the region was subjected to polyphase collisional tectonics. The Scythian Platform subsidence / uplift is a result of

superimposed different subsidence mechanisms: synrift-postrift subsidence, foreland syncompressional subsidence and

subduction-related subsidence of a broad region.

RESUME

Plate-forme Scythe, Caucase et region de la Mer Noire : histoire tectonique et geodvnamique au Mesozoique et

Cenozoi'que.

L’orogenesc est arrivee au stade de collision dans la zone Caucase-plate-forme Scythe a la limite Trias-Jurassique. Durant

la periode comprise entre le Jurassique et PEocene, la region etait situee dans un environnement tectonique d’arriere-arc sujet a

de nombreuses phases d’extension et de compression. A la limite Eocene / Oligocene, le systeme de zone de subduction a ete

transforme en une zone de collision, et durant I’Oligocene-Quaternaire, la region a ete soumise a une tectonique collisionnelle

polyphasee. Les subsidence / soulevement de la plate-forme Scythe resultent de la superposition de differents mecanismes de

subsidence : une subsidence synrift et post-rift, une subsidence syncompression de bassin d*avant-pays et la subsidence d'une

vaste region reliee a la subduction.

NlKISHIN, A.M., CLOETINGH, S.. Brunet. M.-F., Stephenson, R.A., Bolotov, S.N. & Ershov, A.V., 1998.— Scythian

Platform, Caucasus and Black Sea region: Mesozoic-Cenozoic tectonic history and dymanics. In: S. Crasquin-Soleau & E.

Barrier (eds), Peri-Tethys Memoir 3: stratigraphy and evolution of Peri-Tethyan platforms. Mem. Mus. natn. Hist. nat.. 177 :

163-176. Paris ISBN : 2-85653-512-7.

Source: MNHN. Paris

164

ANATOLY M. NIKISHIN ETAL.

INTRODUCTION

The Scythian Platform-Caucasus-Black Sea region has a complicated Mesozoic-Cenozoic history

(Sengor & Yilmaz, 1981; EErcourt et ai, 1993; Kazmin & Sborschikov, 1989; Lordkipanidze,

1980; FlNETTI et al ., 1988; ZIEGLER, 1990; BAZHENOV & BURTMAN, 1990; OKAY et al, 1994;

STAMPFLI & MARCHANT, 1995). We discuss new reconstructions of the regional tectonic history based

mainly on a new scheme of tectonic events in the whole area. We will discuss separately tectonic history

of the Scythian Platform, Great Caucasus, Southern Crimea orogen, Dobrogea, Moesian Platform,

Pontides-Transcaucasus-Alborz magmatic belt and the Black Sea.

MESOZOIC-CENOZOIC HISTORY OF THE SCYTHIAN PLATFORM

The Scythian Platform is located between the Great Caucasus-Southern Crimea orogen to the South,

and the East European Platform to the North. The geological history of the Scythian Platform is

described in NIKISHIN et al ., (this volume), and a summary is presented below (Figs 1-8):

Early-Middle Triassic-Carnian: a major continental rift system originated along the former Late

Paleozoic orogen. Rift related magmatism took place in many rifted basins. Compressional tectonics

took place in pre-Norian time.

Norian-Rhaetian: compressional tectonics, molasse basins origin accompanied by orogenic

magmatism.

Triassic / Jurassic boundary and Hettangian: collisional orogeny.

Sinemurian-early Aalenian: weak rifting in the uplifted Scythian Platform.

Late Aalenian-early Bathonian: regional subsidence of the Scythian Platform, pre-Callovian uplift.

Callovian-mid-Berriasian: a new rift phase took place in the Callovian-Tithonian along the Great

Caucasus-South Crimean belt, in the Dobrogea and the Moesian Platform. The rifting was accompanied

by subsidence of the Scythian Platform.

Berriasian: orogenic event, uplift and thrusting-folding.

Late Berriasian-Barremian: a new system of sedimentary basin configuration originated in the

Scythian Platform; possibly in response to a new weak tensional phase.

Aptian-Albian: a system of rifted basins originated in the Scythian Platform area.

Late Cretaceous: postrift subsidence and tensional events in the Scythian Platform.

Paleocene-Eocene: regional subsidence of the Scythian Platform modulated by stress events.

Eocene / Oligocene boundary ; and Oligocene-early Miocene (Maykop time): rapid syncom-pressional

subsidence of the southern part of the Scythian Platform led to the formation of deep-water basins which

were filled by clay and sandstones.

Middle Miocene-Quaternary: a continent-continent collision has led to collisional tectonics and

orogeny in the Caucasus-South Crimea region, and a few molasse basins originated to the north and

south of the Great Caucasus-Southern-Crimea orogen. North Caucasus Indol-Kuban and Terek-Caspian

molasse basins underwent rapid subsidence phases: 16.5-15, 12.4-9.7, 7-5, 3.7-1.8 and 1.6-0 Ma ago.

Since the middle Miocene a syncompressional bulge has been uplifted along the belt of Karpinsky

Swell-Donets Basin-Ukrainian Shield.

MESOZOIC-CENOZOIC HISTORY OF THE GREAT CAUCASUS

Many different views exist on this issue (MlLANOVSKY & KHAIN, 1963; BELOV & S ATI AN, 1989;

LORDKIPANIDZE, 1980; DERCOURTer al ., 1993; Sharafan, 1995). Below we will discuss a recently

developed working model of the Great Caucasus history:

Triassic / Jurassic boundary: origin of the collisional orogen nearly along the recent Great Caucasus

(Kazmin & Sborschikov, 1989);

Source: MNHN. Paris

MESO-CENOZOIC TECTONIC HISTORY OF SCYTHIAN PLATFORM

165

Sinemurian-early Pliensbachian: the formation of the continental rift-like sedimentary basin (data of

PANOV & Guschin, 1987) in a back-arc environment (?);

Late Pliensbachian-Toar clan-early Aalenian: the development of the deep-water Great Caucasus

trough with rifted, very thin continental crust and basalt volcanism (or local spreading of the oceanic

crust in the Toarcian) (data of Panov & GUSCHIN, 1987; LORDKIPANIDZE, 1980).

Late Aalenian-Bajocian-Bathonian: development of the major subduction-related magmatic belt

along the southern margin of the Great Caucasus trough (LORDKIPANIDZE, 1980; PANOV & GUSCHIN,

1987), inversion deformations inside the Great Caucasus trough;

End of the Bathonian: peak of inversion tectonics, thrusting and folding inside the trough (Panov &

Guschin, 1987; Koronovsky et at.. 1987);

Late Jurassic-Eocene: development of the deep-water flysch trough along the southern part of the

recent Great Caucasus (MlLANOVSKY & KHAIN, 1963; BELOV & SATIAN, 1989); The flysch trough

underwent a complex stress history; for example a tensional phase took place in the Late Jurassic (origin

of the trough); basalt rift-like magmatic and tensional events took place along the trough possibly in

Albian-Cenomanian (LOMIZE, 1969; LORDKIPANIDZE, 1980); during the late Paleocene Great Caucasus

trough underwent a phase of rapid subsidence (possibly due to extension) (BENIJAMOVSKY & SCHERBA,

1995); a weak compressional event took place in the early-middle Berriasian;

Eocene / Oligocene boundary and early Oligocene: compressional tectonics, folding and thrusting at

least along the northern portion of the trough (Milanovsky, 1968; Adamian/ al ., 1989);

Oligocene-early early Miocene: compressional tectonics in the relative deep-water Great Caucasus

trough;

Late early Miocene-Quaternary: polyphase collisional tectonics, formation of the mountains

(possibly since 16.5 Ma ago) and asymmetric fold-thrust belt (MlLANOVSKY, 1968).

MESOZOIC-CENOZOIC HISTORY OF THE SOUTHERN CRIMEA OROGEN

Our model of the geological history of the south-western portion of the orogen is based on a new

interpretation of former and our new data (MURATOV, 1969; Slavin, 1986; Slavin et al ., 1983;

Mazarovich & MlLEEV, 1989 a, b; PANOV et al ., 1994):

Early-Middle Triassic (formations are absent in the folded complex, there are olistoliths in younger

sediments and data on Scythian Platform rifted basins): origin of a rifted basin or passive margin;

Carnian-Norian (data of SLAVIN, 1986 mainly in our interpretation): syncompressional formation of

flysch complex, shallow-water clastic sediments and acidic volcanism;

Rhaetian-Hettangian-early Sinemurian (?) (data of SLAVIN, 1986; PANOV et al ., 1994, mainly in our

interpretation): proposed orogenic phase, folding, uplift followed by formation of shallow-water

carbonate cover;

Late Sinemurian(?)-Pliensbachian-Toarcian-early Aalenian(?): development of the deep-water

trough with turbidite sedimentation in the deepest parts (SLAVIN, 1986; SLAVIN et at ., 1983;

Mazarovich & Mileev, 1989 a; Panov et at ., 1994);

Aalenian-earty Bajocian : intensive orogenic event, thrusting and folding; origin of the Bitak molasse

basin filled by conglomerates (up to 4 km) in the belt between Southern-Crimea orogen and Scythian

Platform (data of MURATOV, 1969; SLAVIN & CHERNOV, 1981 in our interpretation); development of

the Beshuy coal-bearing molasse basin inside the orogen (ROMANOV et al ., 1987);

Late Bajocian-Bathonian: development of the subduction-related volcanic island arc, volcanism

(mainly in the late Bajocian), formation of shallow-water to turbidite sediments (data of Muratov,

1969; Mazarovich & Mileev, 1989 a, b);

Pre-Callovian time : intensive orogenic event, thrusting and folding (data of SLAVIN, 1986;

Muratov, 1969);

Callovian-middle Berriasian: development of the marine basin in the rift-like trough along the

former orogenic belt, filling of the basin by shallow to relative deep water carbonates, conglomerates

166

ANATOLY M. NIK1SHIN ET AL.

and flysch-like deposits (MURATOV, 1969); fold-thrust deformations in the Berriasian (MlLEEV et al.,

1995);

Late Berriasian-Eocene: relatively stable tectonics, covering by shallow-water sediments;

Oligocene-Quaternary: uplift of the area.

MESOZOIC-CENOZOIC HISTORY OF DOBROGEA

North Dobrogea is located between Scythian and Moesian platforms (to the North-East of the

Peceneaga-Camena Fault) and is an uplifted basement of the Scythian platform with the Late Paleozoic

basement (SANDULESCU et al., 1995). The data on the geology of the Dobrogea are summarized in

SANDULESCU et al. (1995) and KRUGLOV & TSYPKO (1988). The following stages of the geological

history in the Mesozoic-Cenozoic can be recognized:

Early Triassic-Camian: continental rifting;

Norian-Rhaetian-Hettangian: postrift subsidence modulated possibly by compressional (?) events;

Sinemurian-Callovian: origin of a relative deep-water rift-like trough with turbidite sedimentation, a

complex history of the trough is not excluded, folding in the pre-Oxfordian time;

Oxfordian-Kimmeridgian: new rift phase, pre-Hauterivian (possibly intra-Berriasian) inversion

tectonics, thrusting ; pre-Late Cretaceous uplift;

Latest late Albian-Campanian: platform subsidence; post-Cretaceous uplift.

MESOZOIC HISTORY OF THE MOESIAN PLATFORM

We will discuss the history of the Varna block located at the eastern part of the Moesian Platform and

the shelf of the Black Sea using data of DACHEV, STANEV & BOKOV in BELOUSSOV & VOLVOVSKY

(1989). Triassic sediments compile a few basins with sediments partly eroded at the Triassic / Jurassic

boundary times. It is possible that the Triassic sediments formed rifted basins or part of a passive

margin. Orogenic deformations and uplift took place at the Triassic / Jurassic boundary. The Early-

Middle Jurassic is presented by shallow-water clastic sediments. The Upper Jurassic-Valanginian

consists of carbonate cover; an uncompensated sedimentation took place in a relatively deep water rifted

basin along the eastern and southern margins of the Varna block. Barremian-Upper Cretaceous forms a

platform cover.

PONTIDES-TRANSCAUCASUS-ALBORZ MAGMATIC BELT

The history of the subduction-related magmatic belts of the Caucasus-Black Sea region is very

complex (DERCOURT et al., 1993; LORDKIPANIDZE, 1980; STAMPFLI & MARCHANT, 1995; OKAY et al.,

1994; ROBINSON et al., 1995). Two Mesozoic-Cenozoic magmatic belts can be recognized in the

Caucasus-Black Sea region: a Triassic magmatic belt with hypothetical position and remnants of the belt

in Northern Iran (STAMPFLI & MARCHANT, 1995), North Caucasus-Turkmenia- North Afghanistan

(Khain, 1979) and possible in the Western-Central Pontides(USTAOMER & ROBERTSON, 1994), and a

Jurassic-Eocene Pontides-Transcaucasus-Alborz magmatic belt (LORDKIPANIDZE, 1980). The history of

the Triassic magmatic belt is not well constrained.

The Pontides-Transcaucasus-Alborz magmatic belt trends at least from the Srednegorie to West

Pontides and East Pontides, and to Transcaucasus- Alborz (LORDKIPANIDZE, 1980). During the Jurassic-

Barremian it is possible to separate a number of events in the magmatic belt history (LORDKIPANIDZE,

1980; TVALCHRELIDZE& Mikhailov, 1985; Sengor & Yilmaz, 1981; ROBINSON etal, 1995). The

climax of volcanism took place in the Transcaucasus during the Bajocian (PANOV & GUSCHIN, 1987;

MlLANOVSKY, 1991).

During the Aptian-Late Cretaceous the magmatic belt was well developed from Bulgaria to northern

Iran. In the Srednegorie (Bulgaria) volcanism initiated in the Campanian (TVALCHRELIDZE &

Source: MNHN. Parts

MESO-CENOZOIC TECTONIC HISTORY OF SCYTHIAN PLATFORM

167

MIKHAILOV, 1985), in the Western Pontides: since the Coniacian (well documented) or even Aptian-

Albian to Campanian or Maastrichtian (GORUR et at., 1993), in the Eastern Pontides - since Turonian to

Maastrichtian (ROBINSON et at ., 1995), in the Transcaucasus belt - since Aptian-Albian to early

Turonian (and locally - to the early Maastrichtian) (TVALCHRELIDZE & MIKHAILOV, 1985;

LORDKIPANIDZE, 1980), in the Alborz - in the late Barremian-Aptian-early Senonian (WENSINK &

Varekamp, 1980).

During Paleogene times subduction-related magmatic belt is known in the Eastern Pontides,

Transcaucasus belt, and in the Alborz (LORDKIPANIDZE, 1980). The volcanism took place in the Eocene

and locally early Oligocene; there are no well documented Paleocene volcanites (LORDKIPANIDZE,

1980; TVALCHRELIDZE & Mikhailov, 1985; Milanovsky, 1991; Robinson et at ., 1995, 1996). The

climax of the volcanism occurred in the mid-Eocene time.

HISTORY OF THE BLACK SEA BASIN

The Black Sea deep-water basin originated as a back-arc extensional structure (ZONENSHAIN & LE

PlCHON, 1986; GORUR, 1988; FlNETTI et at ., 1988). The main problem is a timing of the opening and

geometry of the opening. Western Black Sea Basin originated in the Mid-Late Cretaceous (FlNETTI et

at ., 1988; GORUR et at., 1993), but Eastern Black Sea Basin originated or in Mid-Late Cretaceous

(FlNETTI et at ., 1988; BELOUSSOV & VOLVOVSKY, 1989; OKAY et at ., 1994), or in Paleocene

(Robinson et at ., 1995, 1996), or in Eocene (LORDKIPANIDZE, 1980). We have the following

constraints on the timing of the Black Sea Basin origin based on continental geology:

— during the Aptian-Albian continental rifting took place in Crimea region (Karkinitsky rift system)

and in Western Pontides (GORUR et at ., 1993), weak proposed rift events took place in the Hauterivian-

Barremian in the Crimea too;

— in Bulgaria just to the north of Srednegorie subduction-related volcanic belt rifting took place in

Cenomanian-Turonian (TVALCHRELIDZE & MIKHAILOV, 1985);

— in Georgia to the north of the Transcaucasus subduction-related volcanic belt back-arc late

Turonian-Santonian alkaline basalts are formed (LORDKIPANIDZE, 1980);

— inside the Great Caucasus trough basaltic volcanism of probably back-arc origin took place in

Albian-Cenomanian (LOMIZE, 1969; LORDKIPANIDZE, 1980);

— weak continental rifting took place in Late Cretaceous to the north of Crimea (CHEKUNOV et at .,

1976);

— during the Eocene the Transcaucasus magmatic belt suffered intra-arc rifting with basaltic

volcanism;

— the southern part of Crimea underwent uplift possibly associated with rift shoulder development

during the early-middle Albian; thermal regional subsidence in the Crimea (our data), Western Pontides

(GORUR et at ., 1993) and Dobrogea (SANDULESCU et at 1995) started since the latest late Albian-

Cenomanian;

— inside the Scythian Platform stress changes took place at the Cretaceous / Paleocene boundary;

— the Eastern Pontides were possibly thermally uplifted in Paleocene times (ROBINSON et at ., 1995).

Data on rifting around the Black Sea and on the history of subduction-related volcanic belts show

that the continental rifting in the Western and Eastern Black Sea basins took place in Aptian-Albian

(proposed weak rift event could be in Hauterivian-Barremian); ocean crust spreading took place since

the latest late Albian-Cenomanian to the Late Cretaceous in the Western Black Sea Basin. Eastern Black

Sea Basin suffered at least two stages of rifting: during the Aptian-Late Cretaceous, and, according to

ROBINSON et at. (1995) in the mid-Paleocene.

The Black Sea Basin underwent a number of compressional phases in the Cenozoic. During the

middle Eocene collision of the Rhodope terrane with Moesian Platform and West Black Sea basin took

place (FlNETTI et at ., 1988; BELOUSSOV & VOLVOVSKY, 1989). During the Oligocene-early Miocene

(Maykop time) syncompressional deep-water basins originated at the margins of the East Black Sea

Basin: Sorokin Basin just to the south of Crimea, Tuapse Basin at the boundary between Black Sea and

recent Great Caucasus and Guria Basin to the north of East Pontides (TUGOLESOV et at ., 1985; FlNETTI

168

ANATOLY M. N1KISHIN ETAL.

et al., 1988). During the Neogene-Pleistocene, Black Sea underwent compression and thrusting along

the margins together with molasse basin formation (TUGOLESOV et al., 1985; FlNETTI et al, 1988).

PALEOTECTONIC RECONSTRUCTIONS

Our paleotectonic reconstructions for the region are shown on figures 1 to 8. The geometry of

reconstructions is very simplified. We can recognize the following stages:

Early Triassic-early Carnian: origin of a major continental rift system in the Scythian-Dobrogea-

Moesian region, possibly in a back-arc tectonic environment.

Carnian / Norian boundary-Hettangian: collisional tectonics along the eastern Moesia-Scythian-

Crimea-Great Caucasus-Alborz region. The collision took place between the Eurasia and some

continental terranes (possibly Transcaucasus terrane, Alborz terrane, Western and Eastern Pontides

terranes) (KAZMIN & SBORSCHIKOV, 1989; OKAY et al., 1994; ROBINSON et al., 1995; STAMPFLI &

MARCHANT, 1995).

Sinemurian-Toarcian: the Transcaucasus-Pontides subductional magmatic belt originated (may be

close to position of the former Triassic magmatic belt) possibly in the Sinemurian-Toarcian with north¬

dipping subduction. A major back-arc rifting phase took place and deep-water rifted basin was formed

along the belt Great Caucasus-Southern Crimea-recent margin between Moesian Platform and Western

Black Sea basin-southern margin of the Moesian Platform; rift events took place in the Scythian

Platform and Dobrogea.

Aalenian history: during the Aalenian and at the Aalenian / Bajocian boundary compressional

tectonics took place, orogeny in the Crimea, orogeny and uplift in the Pontides (ROBINSON et al., 1995),

and proposed weak compressional deformations in the Great Caucasus Trough (PANOV & GUSCHIN,

1987).

Bajocian-Bathonian history: during the Bajocian a major subductional magmatic belt was active in

the Transcaucasus-Pontides area (LORDKIPANIDZE, 1980; TVALCHRELIDZE & MIKHAILOV, 1985;

Panov & Guschin, 1987; ROBINSON et al., 1995). Intra-Bajocian inversion tectonics took place along

the Great Caucasus-Southern Crimea basins. This inversion tectonics continued to Bathonian times and

culminated at the end of Bathonian-beginning of the Callovian as an orogeny along the Great Caucasus-

Southern Crimea belt. Regional subsidence of the Scythian Platform and southern half of the Russian

Platform was initiated since the late Aalenian and especially since the late Bajocian. It was

simultaneously with the maximum of the late Bajocian subduction related volcanism in the

Transcaucasus-Pontides belt and inversion tectonics inside the Great Caucasus Trough.

Callovian-mid-Berriasian history: a new rifting phase took place in the Callovian-Tithonian along

the Great Caucasus-Southern Crimea belt, in the Dobrogea area and along the eastern and southern

margin of the Moesian Platform; rift volcanism also occurred in Dobrogea and southern part of the Great

Caucasus. The Transcaucasus-Pontides volcanic belt was active in this time, suggesting a back-arc

origin of the rifting.

Berriasian orogenic event: during the Berriasian time or at the Jurassic / Cretaceous boundary an

orogenic event and associated thrusting took place in the Southern Crimea basin; in the Dobrogea (?),

and in the western Pre-Caucasus area of the Scythian Platform.

Late Berriasian-Barremian history: during this time a new configuration of sedimentary basins was

formed in the Scythian platform; possibly in response to a new weak tensional phase.

Aptian-Late Cretaceous history: major subductional volcanic arcs were active in this time along the

Srednegorie-Westem Pontides-Eastem Pontides-Transcaucasus-Alborz belt. A system of back-arc rifted

basins originated to the North of this volcanic belt: the West Black Sea basin, East Black Sea basin,

extension of the Great Caucasus Trough, and the formation of the rift system in the Northern-Crimea

region. The rifting took place partially along the former magmatic belts: for example remnants of the

same Bajocian volcanic arc exist in the Eastern Pontides and in the southern Crimea. Kinematic

restoration of the Black Sea opening is problematic (FlNETTI et al., 1988; OKAY et al., 1994; DERCOURT

et al, 1993). Constraints are required from restoration of the Caucasus region using precise

paleomagnetic data whereas large scale strike-slip movements cannot be excluded. At present different

Source: MNHN. Paris

MESO-CENOZOIC TECTONIC HISTORY OF SCYTHIAN PLATFORM

169

phases of rifting can be recognized but it is impossible to carry out a kinematic restoration; a number of

rifting phases occurred - hypothetical Hauterivian-Barremian, Aptian-Albian (especially middle-late

Albian), Cenomanian-Turonian, Senonian. Ocean crust spreading took place in the West Black Sea

Basin in the latest late Albian-Cenomanian-Turonian and possibly later. Tensional events took place all

around the Scythian Platform in Mid-Cretaceous time.

Paleocene-Eocene history: in the Paleocene, subductional volcanism was very weak (or absent) in

the Transcaucasus-Pontides belt; a possible rift phase took place in the Eastern Black Sea basin

(ROBINSON et al., 1995). During the Eocene (especially mid-Eocene), Transcaucasus-Eastern Pontides

subductional volcanic belt was very active (LORDKIPANIDZE, 1980); at the time of the maximum

subductional volcanism, a relative rapid platformal subsidence of the Scythian and East European

platforms took place.

Orogeny at the Eocene / Oligocene boundary-early Miocene: at the Eocene / Oligocene boundary the

north-dipping subduction system was changed by collision of the Pontides-Transcaucasus-Alborz area

with Tauride-Anatolide (SENGOR & KIDD, 1979; ADAMIA et al ., 1989; ROBINSON et al., 1995) and

Central-East Iran-Lut terrane (SAIDI, 1995). It led to the uplift of the Transcaucasus region and to

compressional tectonics along the Black Sea-Scythian area.

Middle Miocene-Quaternary orogeny: Arabia / Europe Miocene-Quaternary continent-continent

collision (SENGOR & KIDD, 1979) lead to collision tectonics and orogeny in the Caucasus south Crimea

region, and to syncompressional subsidence of the Black Sea basin.

GEODYNAMICAL HISTORY OF THE SCYTHIAN PLATFORM

The following three main driving mechanisms are proposed for the Scythian Platform subsidence in

the Jurassic-Eocene.

Rifting and postrift thermal subsidence. This model of the subsidence is relatively well constrained as

we have a number of rift stages occurring: after every of the rift phase a postrift subsidence took place.

Loading by a new orogen and compressional stresses. Real regional subsidence of the Scythian

Platform started in the Aalenian and mainly Bajocian simultaneously with orogenic events in the

Caucasus-South Crimean belt. Hence the start of the foreland-type subsidence was connected with the

orogeny.

Subduction-related subsidence of a broad region. The maximum of the volcanic activities of the

Pontides-Transcaucasus-Alborz subduction magmatic belt took place in the Bajocian, Albian-Late

Cretaceous and Eocene. At the same times maximum subsidence of the Scythian Platform and southern

part of the Russian Platform takes place. This supports a causal connections between activity of the

subductional magmatism and subduction itself with a broad platformal subsidence.

During the Oligocene-Neogene-Quaternary, the dynamics of the Scythian Platform was controlled by

collision tectonics. Two main collisional epochs can be separated: Oligocene-early Miocene and middle

Miocene-Quaternary. During the Oligocene-early Miocene two different processes took place:

— compression due to collision with Tauride-Anatolide-Iranian terranes and

— possible roll-back of the subducted lithosphere followed by detachment of the subducted slab.

During the middle Miocene-Quaternary new collisional epoch took place after a short break. The

collision took place between the Europe and Arabia. The collision was irregular with five main phases

of shortening: 16.5-15, 12.4-9.7, 7-5, 3.7-1.8 and 1.6-0 Ma ago. Each compressional event was

accompanied by rapid subsidence of molasse basins followed usually by clinoform sedimentation. The

syncompressional bulge to the north of the Scythian Platform molasse basins suffered uplift since the

early Miocene. The distance between the bulge axis and orogen axis is nearly 350 km. The largest

compressional event took place in the Sarmatian (13.6-9.7 Ma ago): at this time compressional folding

and thrusting occurred inside the Russian Platform up to 1200 km to the north of the Great Caucasus.

Fig. |_g.— Series of paleotectonic maps illustrating the Mesozoic-Cenozoic history of the Scythian-Caucasus-Black Sea region.

FlG. 1-8 — Serie de cartes paleotectoniques illustrant Vhistoire meso-cenozoique de la region plate-forme Scythe-Caucase-Mer

Noire.

Earlk'-Mid TRIASSIC

^-ttUROP

PLATFORM

Korkin

^oesion

- basin

back-arc,

oceanic\ or

basin

52'

Fig. 1.— Early-Mid Triassic. See legend in figure 8.

FlG. 1. — Trias inferieur et moyen. Legende, voir figure8.

Peri-Cospion basin

I—_r-T+y_—_

NogaisK^

\basin N

Kuban^

fc^basin

rans-Caucasus x

X terrane\ x \ x

W.Pontides

\terrane

[TTsubd'.

ethys Ocean

Late TRIASSIC-

-HETTANGIAN

—\Pripyat basin

+ + +

+++.++

per basin

4- + + + +

FToesian

basin

Fig. 2.— Late Triassic-Hettangian. See legend in figure 8.

FlG. 2 .— Trias superieur-Hettangien. Legende, voir figure 8.

Source: MNHN. Paris

Dnieper basin

Voronezh

uucVsus' basin

Cq ^ c C &

^TTT^frrrTTrfm

1 < M ' 1 1111M ' i! 1111 it

Tethys Ocean

— —_— bimbirsk —

_ _ _ Saratov __ :

TN— — basin — —

Early

Mid JURASSIC

+ + + X.— —— .

^ROPEA^IP-tAT^€RW-

—\+ + + \ — — —

Peri-Caspian basin

+ + V— — — — + * ++ \

—X High +

+ + + + + + + + % + +

Ukrainian High + + + + + + + + +

+ + ++ + + + + +/^+ + + + +

a T Kalnnytsky A-_- t_-_ TU RAN

+ + + + + + + + 1y^f + + + + + k basin -/ — — , »— ,

+ + + i + -^\^-_PLATFORM_

+ tr ^^^<^12^^S^ + i-o r »-coocosu2<V_—_

FIG. 3.— Early-Mid Jurassic. See legend in figure 8.

FlG. 3 .— Jurassique inferieur el moyen. Legende' voir figure8.

UROPEAN PLATFORM

onsc

Tethys Ocean

JURASSIC

Late

— — Peri-Caspian basin_—-

TURAN —_

—I&I— PLATFORM

200 Km

Fig. 4.— Late Jurassic. See legend in figure 8.

FlG. 4 .— Jurassique superieur. Legende, voir figure8.

Source: MNHN. Pans

S.Cospian

(Xbasin JX

Tethys Ocean

Mid CRETACEOUS - — /T“\- - Simbirsk

‘— —I Voronezh" Saratov _

+ — basin

—_- Peri-Caspian

EAST EUROPEAN _PLATFORM — — basin “ — — — — —

_—( + + ^s“ Dnieper basin—; :—~——-~-

' — + + _

- —\ukrainian^^C~ ^- —-_(+~+^ — — — —_— J _—_

- —D^jHigh + + +

^vS'b^sin ~-^rAe£f^r SCYTHIAN PLATFORM-_~_-

\ ‘._ ‘LA-

Moesian.p^

basin ->

TURAN PLATFORM

Fig. 5.— Mid-Cretaceous. Sec legend in figure 8.

FlG. 5 .— Cretace moyen. Legende. voir figure 8.

Donets+'

SCYTHIAN PLATFORM

Dagesto

'/basin ,

Rhodope

terrane

Alborz'

Tauride-Anatolide-lronian terranes x x

PALEOCENE-

-EOCENE

Simbirsk

Saratov

basin —

-Pen-Caspian basin

——-EAS^EUROP^ANJ^LATFORM

- Dnieper ba"sTn — — — — —

TURAN PLATFORM

0% _ Peri-Black sea basin—;

<n o

Fig. 6.— Paleocenc-Eocene. See legend in figure 8.

FlG. 6 .— Paleocene-Eocene. Legende. voir figure 8.

Source: MNHN. Paris

MESO-CENOZOIC TECTONIC HISTORY OF SCYTHIAN PLATFORM

173

OLIGOCENE

Peri-Caspian basin

EAST EUROPEAN PLATFORM

basin

Donetsa.

High^

+ Ukrainian High + + +

TURAN

PLATFORM

Peri Black Sea basin

SCYTHIAN PLATFORM

basin

Tauride-Anatolide-Iranian terranes

Fig. 7.— Oligocene. See legencLin figure 8.

Fig. 7.— Oligocene. Legende, voir figure8.

CONCLUSION

Major continental rift system originated in the Early-Mid Triassic to the south of the East-European

Platform. Collision orogeny in the Caucasus-Scythian area took place at the Triassic / Jurassic boundary.

During the Jurassic-Eocene times permanent subduction system was relatively stable, and subduction

related magmatic belt was active along the Pontides-Transcaucasus-Alborz belt; the Caucasus-Scythian

region was in a back-arc tectonic environment with many phases of back-arc extension and

compression; polyphase subsidence complicated by uplift events of the Scythian Platform took place.

At the Eocene / Oligocene boundary the subduction system was transformed into a collision zone,

syncollisional subducted slab roll-back followed by slab detachment took place together with the early

Oligocene wide spread rapid subsidence event in the Scythian-Black Sea area.

During the Neogene - Quaternary the region underwent polyphase collision tectonics.

The Scythian Platform subsidence / uplift is a result of superimposed different subsidence

mechanisms: synrift-postritf subsidence, foreland syncompressional subsidence; subduction-related

subsidence of a broad region and so on.

ACKNOWLEDGEMENTS

The work was funded by the Peri-Tethys Program. The international programmes IGCP-369, INTAS,

EUROPROBE and LITHOSPHERE supported our communications and discussions. Russian Geological

Survey (ROSKOMNEDRA) sponsored our many years field works. We thank E.E. MlLANOVSKY, V.E.

174

ANATOLY M. NIKISHIN ETAL.

Late MIOCENE-

-QUATERNARY

Fig. 8.— Late Miocene-Quaternary. Legende: a. land, mainly plains; b. shallow-water and continental sedimentary basins

mainly; c, relative deep shelf marine sedimentary basin: d. back-arc deep-water basin with oceanic or very thin

continental crust in tensional environment, dominantly; e, remnant deep-water back arc basin; f, continental terrane

accreted to Europe: g. molasse basin; h. intraplate volcanites, mainly basalts; i, subduction related magmatic belt; j, syn-

orogenic intracontinental volcanites; k, rift basin: 1, inverted former rift basin; m. thrust belt; n, strike-slip fault; o,

anticline fold; p. subduction zone; q, passive continental margin; r, areas which underwent syncompressional

deformations and mainly uplift during the Late Triassic-Hettangian and Oligocene; s, areas with syncompressional

deformation in the Aalenian and (or) pre-Callovian times; t, Alpine uplifted orogen. Other symbols are explained on the

maps.

Fig. 8.— Miocene superieur-Quatemaire. Legende : a ter re emergee, principalement des plaines ; b, bassins sedimentaires

d'eau peu profonde et continentaux ; c. bassin sedimentaire marin cotier relativement profond ; d, bassin arriere arc

d'eau profonde avec une croute oceanique ou une croute continental tres fine en environnement extensif dominant; e,

bassin arriere arc residuel d'eau profonde ;f "terrane ” continental accrete a I'Europe ; g, bassin molassique ; h,

volcanisme intraplaque, principalement des basaltes ; i, chaine magmatique liee a une zone de subduction ; j,

volcanisme intracontinental syn-orogenique ; k, bassin de rift; /. bassin de rift inverse ; m, chaine de chevauchement ;

n, faille de decrochement ; o. pli anticlinal ; p, zone de subduction : q, marge continentale passive ; r, zone ay ant subi

des deformations syn-compression et principalement des soulevements durant la periode Trias superieur-Hettangien et

Oligocene ; s, - zone ayant subi des deformations syn-compression a VAalenien et (ou) au pre-Callovien ; t, chaine

alpine soulevee. Les aulres symboles sont expliques sur les cartes.

Khain, E.Yu. Baraboshkin, L.F. Kopaevich, B.P. Nazarevich, D.I. Panov, M.G. Lomize, J.

Dercourt, P. Ziegler, w. Cavazza, V.G. Kazmin, J.P. Cadet, L.E. Ricou, G. Stampfli, E.

Gradinaru, M. Wilson for fruitful discusions. Netherlands Research School of Sedimentary Geology

contribution number 960413.

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Source: MNHN. Paris

Triassic series on the Saharan Platform in Algeria;

Peri-Tethyan onlaps and related structuration

Hamid Art SALEM Sylvie BOURQUIN 2 ', Louis COUREL 121 ,

Berrached Fekirine" 1 , Cherif Hellal 1 ", Leila Mami " 1 &

Mohamed TEFIANI 131

SONATRACH, CRD, avenue du l cl novembre, 35000 Boumerdes. Algerie

,2> Centre des Sciences de la Terre de l’Universite de Bourgogne

UMR CNRS 5561,6, boulevard Gabriel, 21000 Dijon, France

11 Institut des Sciences de la Terre de l'Universite Houari Boum6dienne

BP 32, 16111 El Alia Alger, Algerie

ABSTRACT

The aim of this paper is to understand the pattern of the Triassic deposits over the Algerian Saharan Platform and thus to

establish a basis for structural hypothesis. Review of biostratigraphic and some unpublished palynologic data may open new

perspectives. Palynologic assemblages are described from Ladinian / Carnian boundary in the east and from upper Carnian to

Norian in the north. In the south-eastern part of the Algerian Sahara, new references about Anisian vertebrates remains allow to

date the earliest Triassic series. More generally Triassic deposits commence earlier in the eastern part of the platform. Two

cross-sections in the north of the Saharan Platform were studied, based on high-resolution sequence stratigraphy correlations.

Seven minor stratigraphic cycles recording a complete cycle of base-level fall-to-rise. are observed in these series. They are pan

of the general base-level rise trend of a major stratigraphic cycle, i.e. a retrogradational trend. Main results concern the

diachronism of the boundaries of the traditional Triassic lithostratigraphic units. During the Carnian, the dolomitic sebkha

deposits grade westward, i.e. landward, into fluvial environments. During the Norian-Rhaetian, the evaporite series migrate in

the same direction, while fluvial sandstones step landward in a transgressive pattern. The variations in thickness and facies of

the Triassic series are controlled in details by pre-Triassic paleoreliefs and synsedimentary fracturation. Investigations were

carried out by isopach map analysis, used with circumspection. Isopach plots of the lower series provide evidence of pre-

Triassic paleotopography, mainly in the southern part of the platform. Later reactivation of these fractures was progressively

associated with Triassic NE directions, mainly in the northern part of the platform. In the northern domain, during the Carnian /

Norian stage, NE oriented trough persist, associated with a more regular surface, dipping very gently eastward, toward the

marine peritethyan domain in Tunisia. The E-W elongated S4 salt formation with its First salt deposits above the clastic Triassic

sediments seems to characterize Late Triassic / Jurassic opening of the Atlantic / Alpine Tethys domain.

Ait Salem, H., Bourquin, S.. Courel, L., Fekirine, B., Hellal, C.. Maml L. & Tehanl M., 1998.— Triassic series on

the Saharan Platform in Algeria; Peri-Tethyan onlaps and related structuration. In: S. Crasquin-Soleau & E. Barrier (eds),

Peri-Tethys Memoir 3: stratigraphy and evolution of Peri-Tethyan platforms. Mem. Mus. natn. Hist. nat.. 177 : 177-191. Paris

ISBN: 2-85653-512-7.

Source: MNHN, Pans

178

HAMID AIT SALEM ET AL.

RESUME

Series triasiques sur la plate-forme saharienne en Algerie ; onlaps peri-tethysiens et structuration associee.

Cette note vise a comprendre la geometrie et I'organisation des depots triasiques sur la Plate-forme saharienne en Algerie et

a etablir ainsi une base pour des hypotheses structurales. Une revue biostratigraphique et des donnees palynologiques inedites

ouvrent des perspectives stratigraphiques nouvelles. Des corteges palynologiques sont d6crits, dates de la limite Ladinien /

Camien dans l'Est et du Camien superieur au Norien au Nord. Dans la partie sud-est de la Plate-forme saharienne, des donnees

bibliographiques nouvelles concernant des restes de vertebres permettent de dater de l'Anisien les plus anciennes series

triasiques de la Plate-forme. Plus generalement, les series triasiques sont plus anciennes dans la partie orientale. Deux profils

ont ete etudiSs dans le Nord de la Plate-forme saharienne, a partir de correlations basees sur la stratigraphie s^quentielle haute-

resolution. Sept cycles stratigraphiques mineurs ont ete reconnus dans ces series, repr^sentant un cycle complet de baisse &

montee du niveau de base. Ils s'integrent dans la tendance generale de montee du niveau de base d’un cycle stratigraphique

majeur retrogradant. Ainsi est etabli le diachronisme des limites des unites lithostratigraphiques traditionnelles du Trias

algerien. Au cours du Carnien, les depots dolomitiques de Sebkha gagnent vers l'Ouest, en direction amont, au depens des

environnements fluviatiles. Au cours du Norien-Rhetien, les series evaporitiques migrent dans la meme direction, pendant que

les gres fluviatiles gagnent vers l'amont selon un modele transgressif. Les variations d'epaisseur et de facies des series triasiques

sont controlees dans le detail par les paleo-reliefs ante-triasiques et la fracturation syn-sedimentaire. Ces resultats sont bases sur

une etude critique des cartes d’isopaques. La geometrie des isopaques des series inferieures met clairement en evidence la

paleo-topographie ante-triasique, principalement dans la partie sud de la plate-forme. La reactivation posterieure des fractures

anciennes est progressivement associee a des directions NE, surtout dans le Nord de la plate-forme. Dans le domaine

septentrional, pendant la periode Camien / Norien, une gouttiere persistante d'orientation NE presente un fond plus regulier tres

legerement incline vers l'Est, en direction du domaine peri-tethysien marin en Tunisie. L'allongement E-W de la formation

salifere S4, avec ses premiers depots saliferes au dessus des sediments clastiques triasiques, semble caracteriser l'ouverture

d'age Trias superieur / Jurasique du domaine tethysien atlantique / alpin.

INTRODUCTION

The Mesozoic Saharan Platform in Algeria is a vast Peri-Tethyan epicratonic domain bounded to the

north by the "Saharan Flexure" or "Atlasic Flexure" (BUSSON, 1971). Two types of domain occurred

north of this boundary in Triassic times, northern allochthonous series and southern autochthonous

series (Fig. 1). The allochthonous series of the Maghrebides — the Kabyle Ridge — are part of the

Tethyan domain in palaeogeographical terms (Tefiani et al ., 1991; TEFIANI et al ., 1994; DURAND-

DELGA & TEFIANI, 1994). The autochthonous Triassic series deposited north of the Saharan Flexure are

poorly known in eastern Algeria. A few outcrops along the Tunisian border are related to diapirs (VILA,

1994; PERTHUISOT, 1994) or in the Baborian Zone of Tell (BOURMOUCHE et al ., 1996). In the Saharan

Atlas, because of the depth of burial, there is only seismic data to indicate Triassic series about one

thousand metres thick beneath several kilometres of Mesozoic cover (ViALLY et al ., 1994). Despite the

scarcity of data and analogies with the Algerian (ViALLY et al ., 1994), Moroccan (BEAUCHAMP al .,

1995) and Tunisian Triassic series, the autochthonous north domain is distinguished from the Saharan

Platform by important and rapid thickness variations, in relation with synsedimentary structuration.

The Triassic Saharan Platform stands out as a vast area of relatively thin series (not exceeding 500

m) compared with the supposed thick deposits of the Saharan Atlas tectonic trough. The platform is

open to the east, beyond the Tunisian border, but closed to the south and west where the Triassic

deposits become thinner and even pinch out locally. To the north the Saharan Flexure marks the

boundary of accessible Triassic deposits (in situ Triassic series deeper than 5,000 m). This paper

attempts to highlight the characteristics of clastic and evaporitic sedimentation on the platform and its

possible relations with the eastern marine domain. The significance of structural control will be

revealed.

PERSPECTIVES

Following the pioneering works of Saharan geologists and rare oil industry publications, BUSSON

(1967, 1970, 1971) proposed a stratigraphical, structural and palaeogeographical synthesis of the

Triassic of the Saharan Platform in Algeria which remains a work of reference. Advances have been

made in various fields since.

Biostratigraphical data by ACHAB (1970) and REYRE (1973) have been supplemented by oil industry

reports that unfortunately are unpublished in part. Works by JALIL (1993, 1994), JALIL & TAQUET

Source: MNHN. Paris

TRIASSIC SERIES ON SAHARAN PLATFORM

179

Fig. 1.— Triassic palaeogeographical sketch of Algeria; location map and major structural elements. HR- Hassi R’Mel, HB-

Haoud Berkaoui, HM- Hassi Messaoud.

FlG. /.— Carte paleogeographique schematique du Trias d'Algerie. HR- Hassi R'Mel. HB- Haoud Berkaoui. HM- Hassi

Messaoud .

(1994) and JALIL et al. (1995) on vertebrate remains provide new evidence concerning the central issue

of dating the Zarzaitine Sandstones.

Sedimentological works (BENAMRANE, 1987; HAMEL, 1988; AiT SALEM, 1992; AiT Salem &

HELLAL, 1993; BAYARASSOU, 1994) and oil industry papers (DJARNIA, 1991; JONES & TURNER, 1993;

DANIELS et a !., 1994; Deakins & Daniels, 1994) have attempted to establish correlations between

proximal and distal clastic continental series, transit directions and interpretations of relations with the

marine domain.

Detailed investigations of the variations in facies and thickness of the Triassic series have proved the

control of platform structuration. The influence of the Palaeozoic structural legacy was discussed by AiT

OUALI & NEDJARI (1994) and NEDJARI (1994). Fracturation has been particularly well studied in the oil

reservoir areas (Boudjema, 1987; Hamel, 1988; AiT Salem, 1992; Boudjema et al , 1994;

BAYARASSOU, 1994). Progress can now be made in interpreting the Triassic structuration of the Peri-

Tethyan Platform.

This work focuses on clastic Triassic deposits and covers series older than what is known as the d2

dolomite. This series, with its southern argillaceous lateral equivalents, is so far the first reliable level of

correlation above the pre-Triassic basement on a large part of the Saharan Platform in Algeria. Below

this level BUSSON (1971) already presumed diachronism between elastics and evaporites in some

instances. We seek to open up new perspectives on Peri-Tethyan Triassic onlaps, without claiming to

provide an exhaustive synthesis of the Triassic series. Therefore, for methodological reasons, we shall

attribute no value a priori to long distance correlations between the lithostratigraphical formations

currently used ("Serie inferieure", Tl, T2, etc.). Our aim is to understand the relationships between

proximal and distal elastics and between elastics and evaporites during advances of the Triassic

sedimentation domain over the Saharan Platform.

180

HAMID AIT SALEM ETAL.

REGIONAL FRAMEWORK

The Palaeozoic legacy is clearly marked in the morphology of the base of the Triassic. AiT OUALI &

NEDJARI (1994) and NEDJARI (1994) emphasised the structuration of the Triassic Platform into SW-NE

or SSW-NNE oriented strips in relation with the reactivation of Panafrican and / or Late Palaeozoic

fractures. The occurrence of emerged palaeoreliefs that were gradually buried beneath the base of the

Triassic series has been clearly demonstrated, as for example on the Hassi-R'Mel structure, by BUSSON

(1971). BENAMRANE (1987), BOUDJEMA (1987), HAMEL (1988), Djarnia (1991), BOUDJEMA et al.

(1994) and BAYARASSOU (1994). The same workers, however, emphasised the likely reactivation of

fractures during the deposition of the Triassic of Hassi-R'Mel. Similar conclusions have been drawn for

other sites, especially at Hassi-Messaoud (AiT SALEM, 1992).

Thorough knowledge of structuration of the Saharan Platform in the Triassic would require a

combined study of isopach maps at successive time intervals, a sedimentological analysis across the

entire domain and also a subsidence analysis. At this stage, a preliminary examination of the isopach

maps based on lithostratigraphic correlations can be used to outline the platform structure:

— Westward thinning, inferred from analysis of the "Serie inferieure" Formation isopachs (Fig. 2).

The isopach data are confined to the north of the platform but other data point to the same trend further

south. The "Serie inferieure" clearly thins out on the highs of Hassi-R'Mel, Hassi-Messaoud and

Talemzane.

FIG. 2.— Isopach map of the "Serie inferieure", northern part of the Saharan Platform; isopach values in meters (Hellal,

Fekirine & Arr Salem, Sonatrach doc.)

Fig. 2 .— Carte des isopaques de la Serie inferieure, partie nord de la Plate-forme saharienne ; valeurs des isopaques en

metres (HELLAL, Fekirine & AiT Salem, documents Sonatrach)

Source: MNHN. Paris

TRIASSIC SERIES ON SAHARAN PLATFORM

181

— The distribution and geometry of depocentres and the highs separating them is mapped (Fig. 3)

mainly from detailed studies (well-logs, seismic and gravimetric) of thickness variations at the edges of

the highs. On this scale, N-S alignments predominate in the south whereas SW-NE alignments occur in

the north.

— A large N-S fracture belt crosses the platform from Hassi-Messaoud to El Biod, as reported by

BUSSON (1971), Boudjema (1987), AIT Salem (1992), Bayarassou (1994), Hamel (1988). It

includes highs and fractures. These are essential features in the division of the platform into two

separate areas of sedimentation: the Ghadames Basin to the east and the Oued Mya Basin to the west.

Based on this outline of the regional setting, sequence stratigraphy correlations along two transects

provide new data on facies distribution across the platform during the Triassic. First though it is

essential to define a chronological framework despite the inherent difficulties where continental clastic

series are concerned.

Fig. 3.— Structural sketch of the Saharan Platform under the Triassic cover. After SONATRACH geophysic and drilling data.

FlG. 3 .— Carte structural schematique de la Plate-forme saharienne sous la couverture triasique. D'apres donnees de forages

et geophysiques de la SONATRACH.

BIOSTRATIGRAPHIC DATA AND CHRONOLOGICAL FRAMEWORK

Palynology is the main biostratigraphical tool in what are essentially sandy / silty and argillaceous

series. Works by BlARD (1963), ACHAB (1970) and REYRE (1973) form a starting point. Their

chronological framework was used by BEICIP (1978), Total (1979) and SONATRACH in-house reports

(Mami, 1980; BOURMOUCHE, 1982; MAMI, 1989; Mami & BOURMOUCHE, 1994). Nevertheless

biostratigraphy has still not established a single correlation level. Worse still, the Triassic / Liassic

Source:

182

HAMID AIT SALEM ET AL.

boundary is still not precisely defined by any fossil assemblage or localised in the series. For practical

purposes, since Achab (1970) and BUSSON (1971) the Hettangian d2 marker acts as the ceiling to

supposed Triassic series although the biostratigraphic upper boundary of the Triassic has not been

defined. Below, only the Upper Triassic has been formally identified by palynology. REYRE (1973) did

describe, though, a Disacites assemblage that was not striatiti and was devoid of Classopolis (or with a

few rare individuals) (zone 1) which "might be characteristic of the Middle Triassic plus perhaps part of

the Upper Triassic" in the Gassi-Touil area. In this same eastern region close to the Tunisian border

Busson suggested the presence of a Middle Triassic series primarily on the basis of lithological

correlations with Tunisia.

Unpublished palynological works (Mami, SONATRACH) report the following points:

— At the DAS-1 well (Figs 4, 5) at the Hassi-R'Mel site (2259 - 2262 m) the assemblage of

Enzonalasporites vigens , Tsugaepollenites oriens, Ovalipollis pseudoalatus , disaccates of the genus

Alisporites and the persistence of Circulina are indicative of late Carnian / early Norian age.

— The upper Carnian was recognized in a number of wells by a rich and well preserved assemblage

with Camesporites secatus and the occurrence of Brodispora striata in conjunction with Patinasporites

densus , Vallasporites ignacii , Enzonalasporites vigens , Praecirculina granifer and Duplicisporites

granulatus. This assemblage was collected from wells EBW1 (El Borma region on the Tunisian border,

Fig. 1, 2316 - 2414.50 m), RHA1 (Rhourde el Baguel Trough region, Fig. 1, 3089.90 - 3106.50 m) and

also from wells whose sedimentology is studied below (section Fig. 4): LI4 (2986.50 - 3033.50 m), PG1

(3470 - 3555.90 m) and PHI (3815.70 - 3854 m) on the Talemzane high.

— A third assemblage was found in well EBW1 (2419 - 2504 m) (El Borma region on the Tunisian

border. Fig. 1). The microflora is characterised by the absence of Brodispora striata and persistence of

smaller quantities of Camerosporites secatus and Duplicisporites granulatus. Bisaccates predominate,

especially the genera Pityosporites , Sahnites and Ellipsovelatisporites. Sparse spores are found of the

genera Verrucosisporites and Reticulosisporites. Camerosporites secatus was reported from the Middle

and Upper Triassic by SCHEURING (1970, 1978), SCHURMAN (1977), VlSSHER & KRYSTYN (1978),

VlSSHER & BRUGMAN (1981) and Van DER Eem (1983). The oldest report of C. seccatus comes from

Ladinian (MOY & Traverse, 1986). This third Saharan assemblage suggests an upper-Ladinian /

Carnian age, which are the earliest reported in wells of the Algerian Sahara, and which is consistent with

the findings of REYRE (1993) referred to above.

In view of the location of the assemblages described for the wells, these new palynological data

confirm that the Triassic series commence earlier in the east (EBW1 well). The chronological

framework defined by palynology is, however, too loose to allow correlations between the lower and

upper boundaries of the clastic or evaporitic formations.

Vertebrate remains are known from outcrops of the Zarzaitine series in the extreme south-east of the

Algerian Triassic domain. Since LEHMAN (1957) collections have been increased and datings revised

but the fauna have still not been clearly located in the gritty series. Two assemblages, however, can be

distinguished palaeontologically and chronologically. They are thought to be superposed but in the same

hundred metre high cliff. Work is underway to attempt to resolve the question of their location. The

most recent datings by Jalil (1993, 1994), JALIL & TAQUET (1994) and JALIL et al. (1995) are now

very informative. Remains of Capitosaurids, Trematosaurids and Brachyopoids collected from the cliff

base are dated as early Anisian (Spathyan age is no longer mentioned in the latest publications). Above,

the other assemblage has yielded Aetosaurus and Phytosaurus dated as late Carnian to early Norian

(Jalil et al ., 1995). It is now ascertained that the Middle Triassic occurs in the south-east part of the

Saharan Platform in Algeria. The significance of this finding will be discussed later.

HIGH-RESOLUTION SEQUENCE STRATIGRAPHY OF THE NORTHERN PART

OF THE ALGERIAN SAHARAN PLATFORM: STRATIGRAPHIC CYCLE GEOMETRY

This study was carried out on two cross-sections in the north of the Saharan Platform: an E-W

section between the Tunisian border and the Hassi-R'Mel area (Fig. 4) and a N-S section within the

Hassi-R'Mel area (Fig. 5). All the series are part of the Upper Triassic, Carnian to Hettangian. The

transects are bounded at the top by a dolomitic unit, named d2, which is dated as Hettangian.

Source: MNHN. Paris

TRIASSIC SERIES ON SAHARAN PLATFORM

183

.4 E

Fig. 4.— Triassic east-west stratigraphic cycles geometries between the Tunisian border and the Hassi R'Mel area. Nor: Norian;

I to VII: stratigraphic minor cycles. See figure 5 for the location.

FlG. 4 .— Geometrie des cycles stratigraphiques dans le Trias selon une section est-ouest, de la frontiere tunisienne a la region

d'Hassi R'Mel: Nor: Norien; I a VII: cycles stratigraphiques mineurs. Voir figure 5 pour la localisation.

Fig. 5 — Triassic north-south stratigraphic cycles geometries within the Hassi R’Mel area. See figure 4 for the keys.

Fig. 5.— Geometrie des cycles stratigraphiques dans le Trias selon une section nord-sud, dans la region d'Hassi R'Mel: Voir

figure 4 pour la legende.

Source: MNHN, Paris

184

HAMID AIT SALEM ET AL.

BUSSON (1971) proposed lithostratigraphic correlations and supposed that the sandstone deposits are

more recent at the west of the Saharan Platform and that the upper part of the "Trias greseux" Formation

might be the equivalent of the lower part of the salt deposits in the centre and east of the platform. Only

high-resolution sequence stratigraphy could establish correlations between the basin-centre evaporite

series and the basin-margin clastic series.

Correlations within the Triassic series of the north of the Saharan Platform are based on high-

resolution sequence stratigraphy, where stratigraphic cycles of different scales and durations are

recognised. Correlations and cycle ranks are based on the identification of the smallest stratigraphic

units, i.e. genetic sequences recorded in well logs, calibrated on cores, and their grouping on larger

scales by stacking pattern analysis (CROSS, 1988; HOMEWOOD et al., 1992; CROSS et al ., 1993). Genetic

sequences are defined as packages of strata representing a complete base-level cycle of sediment

accumulation (BUSCH, 1971) with no further details as to base-level cycle origin, cycle duration or cycle

symmetry.

Within the predominantly continental Triassic strata, base-level rise-to-fall turnaround surfaces or

maximum flooding surfaces can be correlated easily across the basin (BOURQUIN & GUILLOCHEAU,

1993; BOURQUIN & GUILLOCHEAU, 1996; COUREL et a /., 1994). In such instances, each genetic

sequence is bounded by base-level rise-to-fall turnaround surfaces and records a complete cycle of base-

level fall and rise. By observing whether the stacking arrangement of genetic sequences are successively

more landward (or deeper or thicker) or more seaward (or shallower or thinner, or more amalgamated)

different scales of stratigraphic cycles are identified.

In the E-W section, the Triassic series are made up of four lithostratigraphic units in the west: the

"Serie inferieure", "Trias greseux", "Argiles inferieures" and "Serie salifere" (Fig. 4). The facies are

more distal to the east and the fluvial sandstones grade laterally eastward into dolomitic sabkha (Fig. 4).

Westward a carbonate unit is more-or-less developed within the "Trias greseux" Formation (Fig. 4).

Within the Hassi-R'Mel area, Triassic series are made up of three lithostratigraphic units (Fig. 5): "Serie

inferieure", "Trias greseux", comprising three units — C, B and A — and the "Argiles inferieures"

(gypseous or anhydritic shales). The "Serie inferieure" Formation and unit C are dated as upper Carnian,

while units. B, A and the "Argiles inferieures" Formation are dated as Norian to Rhaetian (HAMEL,

1988). The three sandstone units start with fluvial deposits overlain by lagoonal dolomitic or anhydritic

shales (HAMEL, 1988). A volcanic episode occurred within this area the Hassi-R'Mel High where

Triassic series are not complete.

Correlations of the two cross-sections (Figs 4, 5) show that the boundaries of the Triassic

lithostratigraphic units are diachronous. Seven stratigraphic minor cycles (I to VII), bounded by

maximum flooding intervals within dolomite shales, are observed within these series (Fig. 4). These

cycles record a complete cycle of base-level fall-to-rise. The turnaround surface from base-level fall-to-

rise on the scale of these cycles cannot be characterized easily either at the base or in the lower part of

fluvial deposits in the absence of core data from all the wells. Within the salt series, each genetic

sequence is made up of halite and shale couples, and the three minor cycles (V, VI, VII) display vertical

evolution where the genetic sequences are increasingly argillaceous. These seven minor cycles are part

of the general base-level rise trend of a major cycle, i.e. a retrogradational trend, that cannot be further

characterized in the study area, exhibiting the transition from fluvial sandstone to halitic sabkha deposits

and then to dolomitic deposits. The biostratigraphic data obtained from some wells (Fig. 4) confirm that

the first four cycles are part of the upper Carnian. However, the lower part of the first cycle could be

older, especially in the eastern wells.

Correlations on the N-S section reveal five stratigraphic minor cycles, bounded by maximum

flooding intervals within lagoonal shales (Fig. 5). On the high, where wells HR9, HR 150 and HR 177 are

correlated, Triassic sedimentation began with the fluvial sandstones of unit B. Two episodes of

volcanism may have occurred in the south of this area (Fig. 5). High-resolution sequence stratigraphy

correlations indicate that the first two cycles of the E-W section are missing in the Hassi-R'Mel area

(Figs 4, 5). Unit C, dated as upper Carnian, is equivalent to the third cycle. Unit B, dated as Norian,

represents the lateral landward equivalent of the base of the "Serie salifere" in the east and can be

correlated with the fifth cycle (first salt cycle). The sixth cycle (second salt cycle) grades westward into

the fluvial deposits of unit A. In the seventh cycle, the halitic deposits migrate westwards into anhydritic

or dolomitic deposits devoid of fluvial sandstones.

These high-resolution sequence stratigraphy correlations establish stratigraphic cycle geometries.

The Triassic series of the north of the Saharan Platform feature seven cycles which are part of a general

Source: MNHN. Paris

TRIASSIC SERIES ON SAHARAN PLATFORM

185

base-level rise trend of a major cycle, i.e. retrogradational trend. Four cycles occurred during the late

Carnian. In the east these cycles are essentially composed of dolomitic sabkha deposits and grade

westward, i.e. landward, into fluvial depositional environments ("Serie inferieure", "Trias greseux",

"Argiles inferieures" and unit C of the Hassi-R'Mel area). During the Norian-Rhaetian, the evaporite

series migrate westward, or landward, while fluvial sandstones of units B and A step landward in a

transgressive pattern. Vertical aggradation predominates towards the top of the salt series, with lateral

transition of the halitic deposits into anhydritic or dolomitic deposits. These high-resolution sequence

stratigraphy correlations could allow correlations in the Saharan Platform between Algerian and

Tunisian Triassic series.

TRIASSIC STRUCTURATION OF THE SAHARAN PLATFORM IN ALGERIA

Reconstructing the structural history of the Saharan Platform in Triassic times in Algeria requires

sound knowledge of the nature and geometry of sedimentary bodies but also involves situating them in

their chronological framework. Biostratigraphic data are unfortunately of little help and sequence

stratigraphy correlations are confined to transects where detailed studies have been undertaken. The

proposed scheme of the distribution of clastic and evaporitic sedimentary series over space and time is

still only an outline but may serve as a basis for structural hypotheses.

First it should be noted that the earliest Triassic deposits on the Mesozoic Saharan Platform are

thought to date from the Middle Triassic and to be located in the Zarzai'tine region, close to the Libyan

border (Fig. 1). Detailed isopach maps by BUSSON (1971) of the region of the south-east of Hamada de

Tinrhert exhibit a series of bands suggesting sedimentation troughs generally open towards the ENE.

The earliest Algerian Triassic deposits are thought to join up with an eastern domain. Marine facies of

the same age are known in Libya and in the Arabian peninsula. Further north, BUSSON (1971) and more

recent works (BOUAZIZ, 1988; BEN ISMAIL, 1991; CHANDOUL et al ., 1993) show that the Triassic series

of southern Tunisia could have been deposited from Scythian times onwards and that Carnian marine

strata are known in the Mekhraneb Dolomite.

In the northern region of the Saharan Platform, BllSSON's proposal (1971) ascribing the greater

thickness of the "Serie greseuse inferieure" in the east to earlier subsidence in the east than in the west is

confirmed by our description of a progressive retrogradation.

The variations in thickness and in facies of the Triassic series are controlled in detail by pre-Triassic

palaeorelief and synsedimentary fracturation. They are investigated by isopach map analyses.

Unfortunately, these are based on lithostratigraphic correlations and not on time-lines. They must

therefore be used with circumspection. They are confined to the northern part of the platform (Fig. 6)

and are designed to highlight structures on the scale of a hundred kilometres within a limited time

interval. This local information must be supplemented by the map that extends further south (Fig. 3).

This map covers, for all the Triassic series, the major features of thickness variations, whose contours

are smoothed to try to depict the main depocentres and highs.

Isopach plots of the lower series (Fig. 2) are the ones most likely to represent diachronous

formations. They provide clear evidence of the pre-Triassic palaeotopography. Whereas the directions

inherited from Panafrican and Hercynian times are clearly N-S in the southern part of the platform, these

orientations which subsist in the northern area bend for the most part to NE, from the Rhourde el Baguel

Ridge to the Hassi-R'Mel Dome via the Dorbane-Dahar Trough, the El Agreb / Hassi-Messaoud Horst

and the Oued Mya Basin (AIT SALEM, 1992). It should be noticed that these orientations mark

extensions of sedimentary units and may include orientations of different fractures or flexuration.

Isopach plots of the formation located between the two sets of volcanic flows (Fig. 7) concern a

poorly defined time interval, given that the flows are probably but not definetly of the same age on the

scale of the same field, but that their ages may vary from field to field. The occurrence of fissure type

volcanism and highly contrasted thickness variations are, however, probable indicators of substantial

fracturation. The NE orientation of the extensions of sedimentary units appears more plainly than before

in a domain where subsidence was probably faster and more irregular.

Isopach plots T1 and T2 on the scale of a vast northern domain (Fig. 8) exhibit two types of

characteristic. South of the region on the map, the transit system via NE oriented troughs persists,

especially in the Oued Mya Basin and the Dorbane Basin. To the north, a much more regular surface

seems to stand out, dipping very gently eastwards towards Tunisia, where subsidence is clearly more

186

HAMID AIT SALEM ETAL.

marked (310 m of sediment on the Tunisian border). By contrast, the series thins out westwards on the

edges of the Talemzane. Hassi-R'Mel, Altai and Hassi-Messaoud Highs.

The S4 Salt Series (Fig. 9) is clearly bounded to the south but its northward extent is unknown. It

shows series of even thickness and thinning of the salt serie towards the west. This thinning results from

halitic facies passing laterally into basin-margin fluviatile sandstones in a retrogradational pattern.

By adding the thicknesses of all the Triassic formations, the lines would mask the stages in the

development of structuration over time. These stages do appear though in the examination of successive

isopach maps, which, however, omit a large southern area of the Saharan Platform. The Mesozoic

structuration is progressively established on the basis of the pre-Triassic legacy in the form of NS-

oriented reactivated fractures and overall palaeotopography. Sedimentary transit orientations bend first

eastward and then NE-oriented fractures are clearly established, in conjunction with fissure type

volcanism in the Hassi-Messaoud and Rhourde el Baguel fields. A vast domain of sedimentation with

series that progressively thicken eastwards, where maximum subsidence occurs, is then established

throughout a large northern region. The salt series towards the east was laid down in a vast E-W trough

bounded abruptly to the south.

Before concluding it should be recalled that the interpretation of the isopach plots is questionable as

thickness cannot be converted directly into subsidence values. Isopach maps merely provide indications

about the geometry of faciological units, which is evidence nevertheless of morpho-structural

parameters. The stress fields responsible for this pattern of development cannot be read directly from the

isopach plots. It is fairly clear though that the WSW-ENE fracturation directions characteristic of the

Source: MNHN, Paris

TRJASSIC SERIES ON SAHARAN PLATFORM

187

Pi G> 7 .— Sopach map of the "Trias eruptif" Formation, northern part of the Saharan Platform; isopach values in meters

(Hellal, Fekirine & Ait Salem, Sonatrach doc.)

Fig. 7.— Carte des isopaques de la formation du Trias eruptif, partie nord de la Plate-forme saharienne ; valeurs des

isopaques en metres (Hellal, Fekirine & A ir Salem, documents Sonatrach)

Atlas Trough and the major Mesozoic fractures of the north of the African Craton appeared

progressively in the course of the Triassic with a NE opening to the Tethys Ocean (BEAUCHAMP et al .,

1995; Jackson et al ., 1995). The fissure eruptions on the Saharan Platform in Algeria are thought to be

associated with differences in subsidence and the development of NE-oriented fractures. The first

volcanic eruptions are thought to date from Camian times when the Triassic fractures of the Algerian

Sahara were initiated. The S4 salt serie with its E-W extension is believed to be representative of

Jurassic structuration.

CONCLUSION

On the Mesozoic Saharan Platform, in view of the vertebrate remains, the earliest middle Triassic

deposits have been located in the Zarzaitine region, close to the Libyan border. In the northern part of

the platform, according to the palynologic assemblages, the Triassic series commence earlier in the east

(Tunisian border): at the Ladinian / Camian boundary, east of the Ghadames Basin. In the west, younger

deposits (upper Carnian) directly lie on the basement. Lower Triassic sediments were deposited in the

north-east trending depocenters only.

According to high resolution sequences stratigraphy correlations, seven minor stratigraphic cycles,

part of the general base-level-rise trend, i.e. retrogradational trend, are observed in these series. In the

lower part, during the Carnian, eastern dolomitic sebkha deposits grade westward, i. e. landward, into

fluvial environments. In the upper part, during the Norian, evaporite series migrate westward or

landward, while fluvial sandstones step landward in an onlapping pattern. Vertical aggradation

predominate towards the top of the salt series, with lateral transition of the halite deposits into anhydrite

188

HAMID AIT SALEM ET AL.

Fig. 8 .— Isopach map of the "T1 and T2" Formations, northern part of the Saharan Platform; isopach values in meters

(Hellal, Fekirine & Ait Salem, Sonatrach doc.)

Fig. 8 .— Carte des isopaques des formations TI et T2, partie nord de la Plate-forme saharienne; valeurs des isopaques en

metres (Hellal, Fekirine & Air Salem documents Sonatrach)

or dolomitic deposits. The description of the progressive retrogradation can be understood by the

development of subsidence area towards the external domain of the platform.

Isopach maps have been analysed across four lithostratigraphic units. Such units are obviously

diachronous but allow to investigate Triassic Saharan Platform structuration. Direction inherited from

Panafrican and Variscan times are clearly N-S in the south. This orientation subsits in the north but bend

for the most part to NE. Contrasted thickness variations, linked with fissure volcanism fracturations, are

good indicators of younger (Carnian) NE directions. North of the Algerian Platform, NE oriented

sedimentation troughs coexist with a much more regular surface in the Ghadames basin, dipping very

gently eastward, toward the marine peritethyan domain in Tunisia. The S4 salt series at the top of the

Triassic series is clearly EW bounded to the south.

The NE opening seems to correspond to the opening of the Neo-Tethys / East Mediterranean system

initiated in Permian time (STAMPFLI et al., 1991). In the Djeffara (South Tunisia) where a continuous

Permian-Triassic marine sequence is known, rifting is at least Permian or older. A thermal contraction

of the thinned crust could have generated the onlap of the Triassic sequence in the Saharan domain. The

Carnian fracturation is found all around the Atlantic-Alpine area and could be related to the Pangea

break-up. This second cycle is superimposed to the thermal subsidence and correspond to the Late

Triassic / Jurassic opening of the Atlantic / Alpine Tethys (STAMPFLI et al ., 1991).

Source: MNHN. Paris

TRIASSIC SERIES ON SAHARAN PLATFORM

189

FiG. 9 .— Isopach map of the "S 6 rie salifere S4” Formation, northern part of the Saharan Platform; isopach values in meters

(Hellal, Fekirine & Ait Salem, Sonatrach doc.)

FlG. 9.— Carte des isopaques de la Serie salifere S4, partie nord de la Plate-forme saharienne ; valeurs des isopaques en

metres (Hellal, Fekirine & A it Salem, documents Sonatrach).

ACKNOWLEDGEMENTS

Critical review provided by G. STAMPFLI improved the manuscript. We gratefully acknowledge

SONATRACH for allowing us to study the subsurface data. We thank also the Peri-Tethys Program for

its financial support. This paper is a contribution to the theme "Enregistrement des phenomenes

biologiques et sedimentaires" of the UMR CNRS 5561 "Paleontologie analytique et Geologie

sedimentaire".

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Mami, L., 1980.— Etude palynologique du sondage REH-1. CRD Sonatrach, Rapport n°162-9-20.

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Perthuisot, V., 1994.— Structures et geometrie des diapirs maghrebins. Essai de synthese. Memoires du Service geologique

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Reyre, Y., 1973.— Palynologie du Mesozoique saharien. Memoires du Museum national d'Histoire naturelle, 27 (C): 1-284.

SCHEURING, B.W., 1970.— Palynologische und palynostratigraphiscITe Untersuchungen des Keupers im Bolchentunnel

(Solothurner Jura). Memoires Suisses de Paleontologie, 88 : 1-119.

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SCHURMAN, W.M.L., 1977.— Aspects of Late Triassic palynology. 2: Palynology of the "Gres et schistes a Avicula contorta"

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New data on the Jurassic and Neogene sedimentation

in the Danakil Horst and Northern Afar Depression,

Eritrea

Mario SAGRI", Ernesto Abbate", Augusto AZZAROLI",

Maria Laura BALESTRIERL", Marco BENVENUTI", Piero BRUNI",

Milvio FAZZUOLI'' 1 , Giovanni FlCCARELLl Marta MarCUCCI 111 ,

Mauro PAPINI", Giulio PAVIA ", Viviana Reale", Lorenzo ROOK"

& Tewelde Medhin TECLE "

(l) Dipartimento Scienze della Terra, Universita di Firenze, Via G. La Pira 4, 50122 Firenze, Italy

<2 ‘ Dipartimento Scienze della Terra, Universita di Torino,Via Accademia delle Scienze 5. 10123 Torino. Italy

(3) Department of Mines. Asmara, Eritrea

ABSTRACT

The Jurassic Antalo Limestone in the Danakil Horst sedimentary cover, the Neogene volcano-sedimentary suite of the Sahel

(Massawa) area and the coeval sedimentary Filling of the northern Afar (Danakil) Depression have been examined. Composite

sections, totalling about 1.100 m, have been measured in the Antalo Limestone resting above the continental Adigrat

Sandstones. Six units have been described and their depositional environments have been related to a carbonate ramp system.

Basinal conditions, marked by dark grey marls and shales, have been recognized in the middle-lower portion of the succession.

Seven depositional sequences, separated by important boundary surfaces, display well-developed system tracts. Ammonites,

foraminifers and calcareous nannofossils allow to assign a late Callovian to early Kimmcridgian age to the studied sections. A

new geological map evidences a complex fault pattern in the northern Danakil Horst. It is likely that, besides normal faulting,

the small Danakil continental block also experienced strike-slip deformation during its detachment from the Nubian plate. In the

Sahel area at the foot of the Eritrean escarpment and above Oligocene Trap Basalts a volcano-sedimentary succession (Dogali

Formation), 550 m thick, consists at its base of lacustrine siliceous deposits interbedded with basalt flows. Fluviodeltaic sands

and gravels prevail in the upper portion with intervening evaporites and patch reef limestones. They are cyclically arranged and

constitute at least four depositional sequences. A proboscidian ( Deinotherium) remain has been found at the base of the Dogali

Formation and indicates an early Miocene age. These syn-rift deposits within the Southern Red Sea rift basin are truncated by

post-rift Boulder Beds fanglomerates. A 500 m thick succession in the upper portion of the Danakil Formation in the Northern

Afar Depression exhibits fluviodeltaic and lacustrine deposits. An early to middle Pleistocene human skull and rich vertebrate

faunas have been found. These fossil findings together with the sedimentologic features of the deposits provide evidence of a

savannah environment prospicient to swamps and shallow lakes. The syn-rift Danakil Formation is uncontormably overlain by

post-rift Boulder Beds. The Sahel and Northern Afar Neogene sediments record the deformational history at the hinge zone

between the Eritrean escarpment and the adjacent subsiding lowlands.

Sagri. M., Abbate, E.. Azzaroli, A.. Balestrieri, M.L.. Benvenuti, M„ Bruni, P.. Fazzuoli, M.. Ficcarelli G..

Marcucci, M., Papini, M.. Pavia, G., Reale. V.. Rook. L. & Tecle, T.M., 1998.— New data on the Jurassic and

sedimentation in the Danakil Horst and Northern Afar Depression, Eritrea. In: S.Crasquin-Soleau & E. Barrier (eds), Peri-

Tethys Memoir 3: stratigraphy and evolution of Peri-Tethyan platforms. Mem. Mus. natn. Hist. nat.. Ill : 193-214. Paris

ISBN : 2-85653-512-7.

Source: MNHN. Paris

194

MARIO SAGRI ETAL.

RESUME

Nouvelles donnees sur la sedimentation du Jurassique et du Neogene dans le Horst du Danakil et dans la depression

du Nord de I'Afar, Erythree.

Trois series sedimcntaires ont ete etudiees : les calcaires jurassiques d'Antalo de la couverture sediment a ire du Horst

Danakil, les formations volcano-sedimentaires neogenes de la region du Sahel (Massawa) et le remplissage sedimentaire de la

partie septentrionale de la depression de I’Afar (Danakil). Des sections composites d'une longueur totale de 1100 m ont 6i6

mesurees dans les calcaires d'Antalo, recouvrant les gres continentaux d'Adigrat. Six unites sont d^crites et leurs conditions de

depots sont misent en relation avec un systeme de rampes carbonatees. Des conditions de depots de bassin, caract6ris£es par des

marnes grises et des shales, ont et£ reconnues dans les parties inferieure et moyenne de la s6rie. Sept sequences de depots,

separees par des limites franches. presentent des 44 system tracts ” bien developpes. L'etude des ammonites, des foraminiferes et

des nannofossiles permettent d'attribuer un age callovien superieur a kimmeridgien aux series etudiees. Une nouvelle

cartographie de la partie septentrionale du Horst Danakil a mis en evidence une tectonique complexe, essenticllement par failles

normales li£es au detachement du bloc Danakil de la plaque nubienne. Cette tectonique distensive est associee a des

mouvements decrochants. Au pied de l'Erythree, dans la region du Sahel, une serie volcano-sedimentaire de 550 m d'epaisseur

se developpe sur les trapps basaltiques oligocenes. Elle est constitute h sa base de depots lacustres siliceux interstratifies avec

des coulees de basalte. Des sables et des graviers fluvio-deltai'ques dominent dans la partie superieure de la serie ou existent

egalement des evaporites et des calcaires recifaux. 11s constituent au moins quatre sequences de depots. Un reste de

proboscibien ( Deinotherium) trouve a la base de la formation de Dogali indique un age Miocene inferieur. Ces depots syn-rifts

du bassin meridional de la Mer Rouge sont tronques par des fanglomerats post-rifts. La serie lacustre et fluvio-deltaique,

epaisse de 500 m, situee dans la partie superieure de la Formation Danakil dans la depression de I'Afar septentrional, a ete

attribute au Pleistocene inferieur et moyen a partir de riches faunes de vertebres et en particulier de restes humains. Ces

fossiles, trouves avec les figures sedimentaires de depot, indiquent un environnement de savane. La formation syn-rift Danakil

est recouverte en discordance par des conglomerats deltaiques post-rifts.

INTRODUCTION

During two field expeditions in Eritrea (November 1994 and December 1995), three geologic

sections through the Danakil Horst and the Northern Afar (Danakil) Depression were surveyed (Fig. 1).

They cross various stratigraphical and structural settings that can be traced on a wide area for which

recent studies are lacking.

We focussed our attention on: (a) the Mesozoic sedimentary cover of the northern portion of the

Danakil Horst; (b) the Miocene to Pleistocene succession near Dogali, west of Massawa; (c) the Plio-

Pleistocene portion of the so-called Red Series near Buia, Northern Afar Depression.

The study area of research theme (a) has been the Danakil Horst, that is a continental block left

behind in the progressive detachment of Arabia from Nubia. It has been rotated by some 30°

counterclockwise (BUREK, 1970; SlCHLER,1980) and is presently bounded on its SW side by the

northern termination of the Afar Depression. Toward NW the latter makes transition into the Southern

Red Sea coastal plain (Sahel) at the foot of the Eritrean escarpment (Fig. 1).

As to the Mesozoic successions of the Danakil Horst, our specific aim was to gain information on

significant events (such as the inception of marine deposition, the age of clastic inflows and of the

anoxic events) that could be correlated with coeval occurrences in Northern Somalia, Ethiopia and

Arabia.

Research lines (b) and (c) cover areas that straddle accross the Southern Red Sea and the Northern

Afar Depression (Fig. 1). Their geological setting is similar since they expose prevailingly clastic.

Neogene successions that can be envisaged as syn- and post-rift deposits within rift basins. The latter are

the Southern Red Sea basin for the Dogali-Massawa area, and the Afar Depression for the Buia region.

The Neogene deposits were investigated in order to provide chronological and paleoenvironmental data

on the different rifting stages in these two transects.

LITHO-, BIO- AND SEQUENCE-STRATIGRAPHY OF THE ANTALO LIMESTONE

The occurrence of the Antalo Limestone in the Danakil Horst has been reported since a long time in

the pioneering works by VlNASSA DE REGNY (1924, 1931). We have been carrying out a detailed study

of this formation in the northern areas close to the Adeilo village (Figs 2, 3) and only a reconnaissance

survey between Amarti and Edd (Fig. 1). Owing to intense tectonic disturbances affecting the Adeilo

area, the Antalo Limestone could not be examined in a single continuous section. Thus, six partial .sec¬

tions have been measured and vertical reconstructions and lateral correlations are presented in figure 3.

Source: MNHN. Paris

JURASSIC AND NEOGENE SEDIMENTATION FROM ERITREA

195

40 °

- 14 °

InOiilillil Neogene sediments

Trap basalts

Amba Aradam Ss

r MmDa Mraaam ^s

Mesozoic H Anta)o Lm & Adigrat Ss.

^ Basement and Paleozoic sediments

Fig. 1.— Simplified geological map of the Danakil Horst, Northern Afar Depression and adjoining Eritrean-Ethiopian Plateau,

with the indication of the study areas (Dogali, Buia. Adeilo). Modified after Kazmin (1975), Merla et al. (1979),

Garland (1980), Drury et al. (1994).

FlG. 1 .— Carte geologique simplifiee du horst Danakil, de la depression de I'Afar septentrional et du Plateau erythreo-

ethiopien, avec la localisation des zones d'etude (Dogali, Buia, Adeilo). Modifie d'apres KAZMIN (1975), Merla et al.

(1979), Garland (1980), Drury et al. ( 1994).

It should be noted that while the transition to the underlying Adigrat Sandstones is frequently found, that

to the overlying Amba Aradam Sandstones is not exposed in the northern part of the Danakil Horst.

According to cursory reconnaissance, the latter transition crops out only in the eastern and southern

areas and is probably marked by marly levels.

Another problem we have been facing is the stratigraphic position of the highly tectonized 50 m-

thick gypsum level which outcrops near Adeilo. Its stratigraphic position surmised in figure 3 derives

from loose regional correlations (south of Amarti, gypsum beds occupy a similar position) and from a

tentative correlation with the Simi Koma key bed (see below).

THE ANTALO LIMESTONE IN THE ADEILO AREA

More than 1,000 m have been measured in detail close to Adeilo and six main rock assemblages

(“ Units ”) (from A to F) have been provisionally distinguished (Fig. 4). Their thickness varies from a

few tens (Unit B) to a few hundreds of meters (Unit C). Each unit is composed of different lithofacies

that have been grouped under different subunits (e.g. Al, A2).

196

MARIO SAGRI ETAL.

PHOTOGEOLOGICAL MAP OF THE

ANTALO LIMESTONE

IN THE ADEILO AREA

(DANAK1L HORST, ERITREA)

Alluvial deposits

Eggerale marls (Unit F)

Eggerale marls and limestones \ t.

(Unit E) \\

Eggerale limestones (Unit D)

Assale mads (Units B-C)

— |— Horizontal and

^ inclined strata

Ade*lo gypsum

Mahur marls and limestones

(Unit A)

Adigrat Sandstones

Basement

Faults

Inferred

faults

Cross

sections

Stratigraphic

sections

5 km

Fig. 2.— Photogeological map and cross sections of the Adeilo area.

Fig. 2.— Carte photogeologique et coupes du secteur d'Adeilo.

Source: MNHN. Paris

JURASSIC AND NEOGENE SEDIMENTATION FROM ERITREA

197

Fig. 3.— Detailed sections of the Antalo Limestone in the Adeilo area. For location see figure 2.

Fig. 3 .— Coupes detaillees des calcaires d'Antalo dans le secteur d'Adeilo (localisation figure 2).

Source: MNHN. Paris

198

MARIO SAGRI ETAL.

Unit A (Mahur and Simi Koma I and II sections. Fig. 3; 0-220 m. Fig. 4). Unit A (220 m thick)

consists predominantly of lithologies varying between blueish limestones and dark grey-blueish marls.

Moreover, it is characterized by the occurrence of more or less frequent bioclastic levels. Shells are

generally broken (coquinas), but also many levels with unbroken shells (lumachellas) are present.

Within marly intervals, coquina levels are generally thin and the shells are completely embedded in a

muddy matrix. If carbonate increases and calcilutitic levels become more frequent, the bioclastic beds

are thicker, coarser, often graded and with an erosional base. In other occurrences, monotypic, well-

sorted and well-cemented lumachellas have been observed. In siliciclastic beds, graded bedding,

horizontal, wavy and hummocky laminations also occur. All these features are consistent with those of

deeper to shallower storm layers (BURCHETTE & Wright, 1992). Unit A can be subdivided in four

subunits (Figs 3, 4). Subunit A1 (0-40 m, Fig. 4) exhibits in its lower part prevailing dark gray, fetid

marly limestones, sometimes nodular and bioturbated, alternating with well-bedded calcilutites and

calcarenites. Bedding joints are generally stylolitic. Bioclasts, in particular pelecypods and brachiopods,

are abundant. In the upper part thick beds of bioclastic calcarenites and, subordinately, of calcilutites

with Hydrozoa are present. They contain cherty nodules and beds, and stylolitized bedding joints.

Fossils: a typical Hydrozoa horizon has been found in this subunit and can be compared with

occurrences at the same stratigraphic level reported in other areas (e.g. Somalia, BRUNI & FAZZUOLI,

1976).The microfauna is rather rich in benthic foraminifers with the classical late Callovian to

Oxfordian assemblage Kumubia palastiniensis Henson, Nautiloculina oolithica Mohler, Lenticulina sp.

and Cyclammina sp. Gastropods, echinoid spines and sponge spicules are also present. The calcareous

nannofossil assemblage is scarce, poorly preserved and not well diversified. It is dominated (Plate 1) by

the Watznaueria genus (W. barnesae (Black), W. britannica (Stradner), W. communis (Reinhardt), W.

fossacincta (Black) and W. manivitae (Bukry)) and includes Cyclagelosphaera margerelii Noel, C.

wiedmannii Reale & Monechi and very rare Lotharingius cf. crucicentralis (Medd) and

Polypodorhabdus spp. C. wiedmannii Reale & Monechi is indicative of a Callovian age (REALE &

MONECHI, 1994). Subunit A2 (40-120 m. Fig. 4) consists of bioclastic calcareous-marly levels

alternating with thinly bedded marly limestones. In the lower part are present calcilutites and numerous

tempestite lumachella beds. Fossils: Peltoceras (Unipeltoceras) sp. indicating a late Callovian age and

numerous brachiopods have been found in the middle portion of this subunit. In the lower part the late

Callovian to Oxfordian assemblage with Kumubia palastiniensis Henson, Nautiloculina oolithica

Mohler, Redmondoides medius (Redmond), R. inflatus (Redmond), Pseudocyclammina sp. and

Glomospira sp. has been recognized. The calcareous nannofossil content is similar to that of Al, but,

because of the greater frequence of marly levels, it is richer and somewhat better preserved. In addition

to those species found in Al, the assemblage includes Biscutum sp., Crepidolithus sp.,

Cyclagelosphaera lacuna Varol & Girgis, Discorhabdus (Palaeopontosphaera) dorse tens is Varol &

Girgis, Podorhabdus grassei Noel, Zeugrhabdotus erectus (Deflandre) and Zeugrhabdotus spp. In the

middle portion of A2 the presence of Stephanolithion bigotii maximum Medd indicates a latest Callovian

/ early Oxfordian age (KAENEL et al ., 1996). Subunit A3 (120-160 m. Fig. 4) is made up mostly by dm-

thick, blueish calcilutitic beds alternating with thinner dark gray to blackish, marly interbeds, often with

Thalassinoides-iype bioturbation and tempestitic lumachella (in particular small ostreids) beds. At the

top of the section thick beds of bioclastic calcarenites appear. Fossils: particularly toward the top, this

subunit is relatively rich in sponge spicules and subordinately in brachiopods. Redmondoides medius

(Redmond) and Bathysiphon sp. have been determined. Biomicrites at the top of this subunit contain

Kumubia palastiniensis Henson and Nautiloculina oolithica Mohler, that are indicative of a late

Callovian to Oxfordian age. In the lower half of A3 the calcareous nannofossil assemblage is again

scarce, poorly preserved, not diversified and dominated by Watznaueria as in Al. The disappearance of

Cyclagelosphaera wiedmannii Reale & Monechi possibly indicates an Oxfordian age. In the upper half

the nannofossils are almost absent and just few Watznaueria persist. Very rare and questionable

Mitrolithus sp. are also present. In the lower part of subunit A4 (160-220 m, Fig. 4) a regular alternance

of dm-thick beds of more or less marly limestone, often including lumachella beds and of marls,

sometimes nodular, is present. In the upper part lumachella beds are gradually replaced by thin siltstone

beds and then by thicker calcareous terrigenous sandstones. The upper part of this subunit, organized in

thickening- and coarsening-up cycles, consists of dm-thick beds of calcareous sandstones with

horizontal, cross and hummocky lamination and thin silty-marly interbeds, passing upwards to a 5 m-

thick coarse terrigenous calcareous sandy bed-set (the Simi Koma key bed). The latter includes cm- to

dm-sized calcarenite clasts and nodules and lenses of blackish chert. Horizontal, cross and wavy

Source: MNHN. Paris

JURASSIC AND NEOGENE SEDIMENTATION FROM ERITREA

199

Marls and

shaly marls

Limestones and

marly limestones

Terrigenous sandstones

HST - Highstand system tract

TST - Transgressive system tract

SMF - Surface of maximum flooding

SB - Sequence boundary

Inner ramp (IR)

Mid ramp (MR)

Outer ramp (OR)

Basin (B)

Fig. 4.— Composite section summarizing the logs of Fig. 3, with the environment and the sequence stratigraphy interpretation.

Units and sub-units as in figure 3.

Fig. 4 — Coupe synthetique des logs de la figure 3. avec les paleoenvironnements et 1’interpretation en termes de sequences

stratigraphiques. Mime unites et sous-unites que figure 3.

Source: MNHN, Paris

200

MARIO SAGRI ETAL.

laminations are well evident as well as a thickening and coarsening up pattern that we relate to the

progradation of a silicoclastic deltaic body. Fossils: foraminiferal assemblages are rather poor in this

subunit, with the exception of some levels particularly rich in Redmondoides sp. together with

Pseudocyclammina sp. and Ammodiscidae. The calcareous nannofossil assemblage is present only in the

lower portion, although it is scarce and poorly preserved with very rare Cyclagelosphaera margerelii

Noel, C. lacuna Varol & Girgis, Watznaueria barnesae (Black) and Polypodorhabdus spp. Age of Unit

A: the above mentioned ammonites and the calcareous nannofossils content indicate a late Callovian age

from the middle part of subunit A2 downward, and an early Oxfordian age upward. The foraminiferal

assemblages are consistent with these datings.

Unit B (Simi Koma II and Assale sections, Fig. 3; 220-280 m. Fig. 4). In its lower and middle parts

dm-thick beds of blueish calcilutites alternate regularly with more or less silty dark grey marls,

tempestitic lumachella beds and with rare thin beds of calcareous, hummocky-laminated, sandstones. In

the upper part the storm layers become rare and eventually disappear. The thickness of the intercalated

marly levels increases, whereas the thickness and the frequency of the calcilutitic beds decrease.

Calcilutites and marls with their dark colour and abundant small sulphide crystals indicate hypoxic to

anoxic conditions of the sedimentary environment. These conditions are present also in the lower part of

the overlying subunit Cl. Fossils: fragments of pelecypods, brachiopods, benthic foraminifers and

sponge spicules are often present. Lenticulina sp. occurs at the base of the Assale section. As to the

calcareous nannofossils, they are absent in the lowermost portion of B of the Simi Koma section, but

they appear upward in the succession with Crepidolithus sp., Cyclagelosphaera margerelii Noel,

Lotharingius hauffii Grim & Zweili, Watznaueria barnesae (Black), W. britannica (Stradner), W.

communis (Reinhardt). The assemblage is indicative of a generic Middle to Late Jurassic age. Age: the

age of Unit B is possibly early to middle Oxfordian based on its stratigraphic position.

Unit C (Assale and Simi Koma II sections, Fig. 3; 280-520 m. Fig. 4). Subunit Cl (280-360 m, Fig.

4) is composed by 10-20 cm thick beds of dark grey, sulphide-rich, calcilutites and dark grey marly,

often yellowish, interbeds. Subunit C2 (360-400 m. Fig. 4) consists of yellowish-grey, splintery marls

with rare levels of marly limestones. In subunit C3 (400-520 m. Fig. 4) the calcareous content increases

and levels, that are prevailingly marly, alternate with calcareous marls. They are both yellowish-pink.

Fossils: Cl yielded some Perisphinctes sp. and Epimayaites sp., while Perisphinctes sp. was found at

the base of C3. Subunit C2 contains abundant ammonites. The most significant are Epimayaites

falcoides Spath, " Perisphinctes " orientalis Siemiradzki and "Perisphinctes" aff. subcolubrinus

Siemiradzki that indicate the middle / late Oxfordian boundary. Pelecypods, ostracods and echinoderms

often appear as bioclasts. In the C Unit, foraminifers are scarce and represented only by some

Redmondoides medius (Redmond), R. inflatus (Redmond), Spirillinidae and Eponides sp. Particularly

in the middle portion of Unit C, the calcareous nannofossils are quite abundant and diversified

and moderately preserved. Their assemblage is dominated by the Watznaueriaceae (Watznaueria

barnesae (Black), W. britannica (Stradner), W. communis (Reinhardt), W. contracta (Bown &

Cooper), W. manivitae (Bukry), and less abundant Cyclagelosphaera margerelii Noel, C. lacuna Varol

& Girgis, Lotharingius cf. crucicentralis Medd, L velatus Bown & Cooper, L. hauffii Grim &

Zweili) and includes, among others. Biscutum dubium (Noel), Crepidolithus sp., Discorhabdus

( Palaeopontosphaera ) dorsetensis Varol & Girgis, Polypodorhabdus spp., Stephanolithion bigotii

bigotii Deflandre, Zeugrhabdotus erectus (Deflandre) and Zeugrhabdotus spp. The presence of

Stephanolithion bigotii bigotii Deflandre in the middle portion of Unit C indicates an age younger than

early Callovian. Age: the age of Unit C is middle to late Oxfordian based on ammonites.

Unit D (Eggerale section. Fig. 3; 520-700 m, Fig. 4). In the lower part (Dl: 520-580 m, Fig. 4) blue-

greyish calcilutites are prevailing. They are fetid, sulphide-rich, dm-thick beds, with rare and thin marly

interbeds. In the field they constitute a prominent cliff. The middle part (D2: 580-620 m, Fig. 4) consists

of thickening- and coarsening-up cycles of calcilutites and calcarenites in beds. 30-60 cm thick

alternating with marls. In the upper part (D3: 620-700 m. Fig. 4) calcilutitic beds, 30-60 cm thick, are

present and form a second steep cliff. Upwards, the bed thickness decreases and there are nodules and

lenses of chert and marly intercalations. Fossils: sponge spicules and late Callovian to Oxfordian benthic

foraminifers (Kurnubia palastiniensis Henson, Nautiloculina oolithica Mohler, Ammodiscidae and

Lituolidae) and rare radiolarians are present.The calcareous nannofossil assemblage, indicative of a

Middle to Late Jurassic age, is scarce, poorly preserved and not diversified. The genus Watznaueria is

Source: MNHN. Paris

JURASSIC AND NEOGENE SEDIMENTATION FROM ERITREA

201

still the most abundant with W. barnesae (Black), W. britannica (Stradner), W. communis (Reinhardt)

and W. manivitae (Bukry). Very rare Lotharingius velatus Bown & Cooper, L. hauffii (Griin & Zweili)

and Cyclagelosphaera margerelii Noel are also present. Age: the age of Unit D is possibly late

Oxfordian - early Kimmeridgian, based on its stratigraphic position.

Unit E (Eggerale section, Fig. 3; 700-960 m, Fig. 4). Marly limestones and marls alternate with thin

bed-sets of calcilutites and bioclastic calcarenites. The marly limestones are often bioclastic and

lumachella beds are present as well. The marls show yellowish or pink-violet colour and evident fixility

or nodularity. Calcilutites are dark grey, sometimes yellowish or reddish, due to oxidation of small

sulfide crystals. Bioturbation is often present. Fossils: bioclasts of pelecypods, gastropods, echinoderms

and rare radiolarians, benthic foraminifers, encrusting foraminifers, serpulids, ostracods occur. As to the

foraminifers, Alveosepta jaccardi (Schrodt), Riyadhella sp., Lenticulina sp. and Ammodiscidae are

present. Saccocoma sp., sponge spicules and crinoidal spines have been also found. In both units E and

F almost all the samples are barren of calcareous nannofossils. Only sporadic Discorhabdus sp. and

Watznaueria sp. have been recognized. Age: late Oxfordian - early Kimmeridgian on the base of the

Alveosepta jaccardi occurrence.

Unit F (Eggerale section, Fig. 3; 960-1080 m. Fig. 4). Subunit FI (960-1020 m, Fig. 4) consists of

often reddish calcilutites in dm-thick beds and subordinate bioclastic calcarenites with thin marly

interbeds. Marly levels and sets of dm-thick calcilutitic beds characterize subunit F2 (1020-1080 m. Fig.

4); in its upper part bioclastic calcarenites and dm-thick tempestitic beds with fragments of pelecypods

and gastropods are frequent. All calcarenite samples contain abundant silt- to sand-size quartz grains.

Fossils: the poor foraminiferal assemblage includes Alveosepta jaccardi (Schrodt), Riyadhella sp.,

Paraurgonina sp. and Ataxophragmiidae. Age: late Oxfordian - early Kimmeridgian on the base of

Alveosepta jaccardi (Schrodt) occurrence.

ENVIRONMENTAL INTERPRETATION AND SEQUENCE STRATIGRAPHY

In the Adeilo area the Antalo Limestone is prevailingly made up by calcareous and marly beds

arranged in 1 to 10 m thick asymmetric cycles. These cycles, that are exclusively subtidal, exhibit the

following attributes: a- within each cycle shaly-marly beds are prevailing at the base and calcareous

beds at the top; b- tempestite beds and / or calcarenites are present in the upper portion of several cycles;

c- the thickness of the calcareous beds increases upward whilst that of the marly beds decreases. These

features indicate upward shallowing cycles. Since the base of each cycle records an increase in water

depth, the cycles can be envisaged as parasequences (Van WAGONER et ai , 1988).

These features support the inference that the Antalo Limestone was deposited in a shaly and

calcareous homoclinal ramp (sensa BURCHETTE & Wright, 1992). Since slumpings and breccia bodies

as well as evidence of synsedimentary faulting are lacking, the floor of the ramp should have been very

gentle. This is suggested also by the extensive areal occurrence of sedimentary units with the same

characteristics.

The variability of the lithological and sedimentological features of the parasequences led to recognize

four sub-environments such as those that appear in the schematic inner ramp to basin section of figure 5.

These sub-environments are: a- an inner ramp with prevailing high-energy, carbonate, bioclastic sandy

barrier; b- a mid ramp with frequent calcareous bioclastic storm layers, alternating with marly and shaly

beds; c- an outer ramp with rare storm reworking, abundance of shaly and marly levels, and d- a basin

with shaly-marly sediments and pelagic faunas (radiolarians, ammonites).

We can refer our subunits in the Adeilo succession to these sub-environments (Fig. 4). Starting from

the bottom, subunit Al is a shoaling up sequence deposited at first on a mid ramp with hypoxic

conditions and then on an inner ramp with medium to high turbulence. Features in subunit A2 indicate a

deepening toward an outer ramp environment, followed by fluctuations within a mid ramp. The

environment in subunit A3 exhibits again a shallowing up evolution from a mid ramp to a restricted

inner ramp. Subunit A4 can be referred, at first, to a mid ramp shallowing upward to an inner ramp, and,

according to the thickening- and the coarsening-up pattern, to the progradation of a siliciclastic delta

(Simi Koma key bed). During the sedimentation of Unit B a progressive deepening occurred, although

limited to a mid ramp environment.

In subunit Cl an initial outer ramp gave place to mid ramp; in C2 basinal conditions, marked by

202

MARIO SAGRI ETAL.

3 EH SSI 53 ESI EEB 8

1 2 3 4 5 6 7

Fig. 5.— Lithology of parasequences and different sub-environments in a schematic inner ramp to basin cross-section. Legend:

1, shales; 2, marls; 3, marly limestones; 4. tempestitic coquina beds; 5, calcarenites; 6. calcilutites; 7, ammonites, a,

mean sea level; b. fair-weather wave base; c, storm wave base.

FlG. 5 .— Lithologie des parasequences et des sub-environements dans une rampe interne schematique. I, shales ; 2, marnes ;

3, marno-calcaires ; 4, calcaires coquillers ; 5, calcarenites ; 6, calcilutites ; 7, amnwnites, a, principaux niveaux

marins : b, limite d'action des vagues ; c, limite d'action des vagues de tempetes.

marly lithologies and extensive occurrences of ammonites, prevailed. Later on, conditions changed back

to mid ramp through some fluctuations (C3).

As to unit D, the environment varies from an inner, restricted ramp in the lower part (Dl) to a mid-

inner ramp in the median part (D2), and, finally, again to an inner ramp (D3). A mid ramp environment

with a relative deepening in the median portion is recognizable in unit E, and a variation from inner

ramp to mid ramp in unit F.

A cyclic repetition of sub-environments is thus evident in the succession that also includes larger-

scale cycles framed between major surfaces (SB 1 to 7, Fig. 4). In the field these surfaces correspond to

morphological contrasts since they are placed where there is an abrupt change between more calcareous

bed packets and the overlying shaly / marly portions. Accordingly, the depositional environment shifts

upsection toward a relevant deepening. These surfaces, sometimes with unconformity features, comprise

bed packages from several tenth to several hundred meters thick and cover a time span of one up to

some million years. They can be regarded as sequence boundaries framing depositional sequences,

possibly of the third order.

From bottom to top, the oldest surface (SB1) corresponds to the boundary surface between the

Adigrat Sandstones and the Antalo Limestone. The second surface (SB2) occurs at the top of subunit

Al, above the level with Hydrozoa and corals. The third surface (SB3) is placed on the top of the Simi

Koma thick arenaceous bed packet, the deltaic body probably connected laterally with the evaporites of

the Adeilo section. The fourth surface (SB4) is present at top of subunit Cl, and the fifth surface (SB5)

on the top of the lower limestone cliff corrensponding to the subunit DL The sixth surface is placed at

the top of the upper limestone cliff corresponding to the subunit D3, and the seventh surface (SB7) at the

top of calcareous subunit FI.

According to this interpretation, these surfaces bound six depositional sequences, namely S1 (40 m

thick), S2 (170 m thick), S3 (130 m thick), S4 (230 m thick), S5 (120 m thick) and S6 (330 m thick)

(Fig. 4). These sequences show more or less developed intervals with a retrogradational arrangement of

parasequences, that are considered as transgressive system tracts, and generally thicker intervals with

progradational arrangement of parasequences (highstand system tract).

It is to be noted that in the sequence S4 the whole transgressive system tract corresponds to the

bacinal C2 subunit and its surface of maximum flooding corresponds to a large-scale relative deepening.

Unfortunately, available chronostratigraphic data are not adequate for a precise dating of the

sequence boundary surfaces and do not allow a close comparison with the global sequence boundaries

(Vail etal ., 1984; Hallam, 1988).

Source: MNHN. Paris

JURASSIC AND NEOGENE SEDIMENTATION FROM ERITREA

203

CHRONOSTRATIGRAPHIC CONSIDERATIONS AND CORRELATIONS

WITH ADJOINING REGIONS IN EAST AFRICA

In the study area the Antalo Limestone was deposited in a time span extending from late Callovian to

early Kimmeridgian. Some comments can be added to these age attributions. As to the lower age limit,

that is the invasion of the Antalo sea above the terrestrial to transitional Adigrat Sandstones, we can

observe that this transgression is widespread in the whole East Africa as well as in Arabia (the so called

Eligmus fauna of ARKELL, 1951). For its large areal extension we can assume that it was connected with

an outstanding eustatic rise. During the Callovian there is an important phase of coastal onlap at 155 Ma

(HAQ et al., 1988). Even if the datings of the lower and upper boundaries of the Callovian are uncertain

(see Harland et al ., 1982; ODIN & ODIN, 1990), we are inclined to relate the basal flooding surface of

the Antalo Limestone in Adeilo to this 155 Ma old rise because of the Peltoceras (uppermost Callovian,

ca. 152 Ma) occurrence 80 m above the base of the studied succession.

It is noteworthy that the greatest deepening of the basin was reached at ca. 148 Ma (middle / late

Oxfordian boundary) as suggested by the marls with ammonites of Unit C, 400 m above the base.

The data obtained from the Antalo Limestone in the northern Danakil Horst can be matched with

those from Ethiopia, Somalia and NE Kenya. We have seen that the inception of marine sedimentation

occurred in the late Callovian in the Adeilo area and we assumed that this event can be framed in the

large-scale marine transgression in the whole East Africa. Some well-known sections, such as Sa Wer

(ABBATE et al., 1974) and Bihendula (BUSCAGLIONE & FAZZUOLI, 1987; ABBATE et al., 1994) in

Somalia, the Blue Nile gorge (FlCCARELLL 1968; RUSSO et al., 1994), Mekele (BOSELLINI et al., 1995)

and Dire Dawa (FlCCARELLl et al., 1975) in Ethiopia can be taken for reference. It has to be added that

some local differences (from late Bathonian to early Oxfordian) exist in dating the first marine episodes,

but they can be imputed to a heterochronous transgression or, more likely, to poor biostratigraphic

calibrations. However, a well documented precocious transgression occurred during the early Toarcian

(SOEC, 1954; CANUTI et al., 1983; BOSELLINI, 1989; LUGERer al., 1990) limitedly to some areas (e.g.

Ahl Mado in northern Somalia and Lugh-Mandera basins in^southem Somalia / northeastern Kenya).

A closer comparison between the Antalo Limestone at Adeilo and in other East Africa successions

shows that the late Callovian Hydrozoa limestones at the top of subunit Al (Mahur section, Fig. 3) can

be correlated with the lithologically similar and possibly coeval level at the top of the Bihen Limestones

at Bihendula (BRUNI & FAZZUOLI, 1976).

Further correlations can be traced among levels showing remarkable deepenings of the basin. The

middle / upper Oxfordian marls with ammonites of Unit C of the Adeilo area correspond to coeval and

similar basinal deposits with belemnites and / or ammonites, that are common in East Africa (see

Warandab depositional sequence of BOSELLINI, 1989).

At a very short distance from the Adeilo area, the Mekele outlier in Tigrai (MERLA & MlNUCCI,

1938; ARKIN et al, 1971; BEITH, 1972; BOSELLINI et al., 1995) share some common features with the

succession described in this paper. In both areas marly and calcareous lithotypes are arranged in

parasequences with frequent storm layers denoting a ramp environment. Moreover, a thick marly

interval with ammonites occurs in the middle portion of both sections, but siliciclastic and / or evaporitic

intercalations are lacking in the lower to middle portion of the Mekele section. Interestingly, the total

thickness of the Antalo Limestone increases from the Mekele area to the Danakil Depression.

STRUCTURAL COMMENTS ON THE ADEILO AREA

A photogeological map at 1: 100,000 scale (Fig. 2) with several field controls is presented for the

Adeilo area. It was meant to trace lateral and vertical extension of the sedimentary units distinguished in

the Antalo Limestone and to provide information on the structure of this area for which detailed

mapping is still lacking. The photogeologic study, together with field observations, testifies at least for

the Adeilo area, a quite complex structural setting. On the basis of reconnaissance work a similar

complexity seems to be common also in the eastern and southeastern slopes of the Danakil block (see

MOHR, 1967). In the mapped area different structural situations can be distinguished.

204

MARIO SAGRI ETAL.

The northern side shows a homoclinal structure with bedding slightly (max. 15°) tilted toward north

and. limitedly to the Mt. Eggerale area, toward NW. This structure is affected by syntethic E-W trending

normal faults with northern downthrown blocks. An important NW-SE trending family of faults has a

possible left transcurrent component suggested by block offsets, rectilinear traces of regional extension

and a few observed striated mesofaults. Large knee folds are locally present (see sect. V, Fig. 2).

The Mt. Assale topographic high is a large downfaulted block bounded to the east by a NNE-SSW

fault system and to the SW by a NW-SE fault system. Only in proximity of the main fault systems the

dipping becomes variable and steeper. The structure is further complicated by a normal fault system

with an E-W trend.

SYN- AND POST- RIFT DEPOSITS IN THE SAHEL AND IN THE NORTHERN AFAR REGIONS

The syn- and post-rift Neogene deposits of the Sahel coastal plain and Northern Afar Depression

record important tectonic events that affected these zones and the surrounding regions. They outcrop on

both sides of the Afar Depression and in the hills west of Massawa (Fig. 1) above Tertiary Trap basalts,

Mesozoic sediments and Precambrian basement (DAINELLI, 1943; Frazier, 1970).

The succession comprises predominantly continental deposits with an exposed thickness of several

hundreds of meters (DAINELLI, 1943; BANNERT etal, 1970; Kazmin, 1975; GARLAND, 1980).

The sedimentary units recognized are generally bounded by sharp erosional surfaces, paleosols and

angular unconformities.

THE SAHEL AREA

In this area a 650 m thick volcano-sedimentary succession (Dogali Formation) rests above Oligocene

(28 Ma, DRURY el al., 1994) Trap basalts and is unconformably overlain by Desset Boulder Beds

(Porro, 1935; Kazmin, 1975) (Fig. 7).

The Dogali and Desset formations merge laterally offshore into the shallow marine, up to 7,000 m

thick, sediments of the Habab, Dunishub and Amber salt formations (Carella & SCARPA, 1962;

Sestini, 1965; Kazmin, 1975).

The dogali formation

Composite sections, totalling 500 m, have been measured in the Dogali, Desset and Gabeichelti areas

(Fig. 7). They show that the lower portion of this formation (lower Dogali Formation) is predominantly

siliceous, while the upper one (upper Dogali Formation) consists mainly of pebbly sands,

conglomerates, reefal limestones and gypsum beds. More in particular, the lower portion is made up by

siliceous, fine grained, yellow, brown and dark grey laminated or nodular clays and silts in 0.2-0.5 m

thick beds with rusty hematitic coats. The silty beds are graded and contain root marks and leaf prints.

Volcanoclastic green sandy beds, 0.2-1.5 m thick, massive or horizontally laminated are interbedded in

the siliceous lacustrine deposits. They are graded and contain reed fragments, leaf prints and clay clasts.

These lithofacies may be interpreted as deposited in fresh-water ponds during volcanic quiescence with

reduced clastic inflows.

A lenticular, 0.5-1.5 m thick level of brown massive or cross-stratified pebbly sandstones is also

present toward the base of this formation. It contains well-rounded pebbles derived from Trap basalts

and associated acidic volcanites, silicified tree trunks, mammalian bones among which a newly

discovered Deinotherium remain (a fragment of lower jaw with teeth).

The lower Dogali Formation with a total thickness of 350 m consists in its upper part of 200 m of

basalts.

The upper Dogali Formation shows diversified lithologies that are cyclically:

— gravelly facies. Massive, normal and reverse graded pebbles with abundant reddish sandy matrix.

Clasts are rounded, imbricated and polymodal in size. Horizontal and trough cross lamination are

common features of the pebbly beds that are 1-2 m thick. Locally, they are amalgamated and lie on

erosional surfaces. Clasts up to 40 cm are polymictic and derive from basement rocks, basalts and acidic

Source: MNHN. Paris

JURASSIC AND NEOGENE SEDIMENTATION FROM ERITREA

205

o o

Terraced alluvial cobbles

O o

and sands (Plio-Pleistocenel

Saati

DogaJi

Desset

Hubbet

Trap basalt flows and dykes

(Late Oligocene-Miocene)

Basement (Precambrian)

0 1 2km

fault

V* '*

evaporitic

*-•-7 key bed

SCHEMATIC GEOLOGIC MAP OF THE

DOGAU AREA, MASSAWA

ERITREA

Alluvial, fluvial, aeolian

deposits (Quaternary)

transitional and shallow marine

deposits, sands, silts, reefal

limest., evaporites (Miocene?)

lacustrine siliceous clays,

silts and interbedded

basalt flows (Miocene?)

Fig. 6.— Sketch map of the Dogali area (west of Massawa). based on field survey and photogeological interpretation.

Fig. 6.— Carte schematique de la region de Dogali (a I’ouest de Massawa). d'apres des etudes de terrain et l'interpretation

des photographies aeriennes.

Source:

Dogali Fm. Desset Fm.

206

MARIO SAGRI El'AL.

150 -

100 -

v v v

v v v vl

550

p 500

yvfe'j

450

V V V

V V V V

V V V

v v v v

V V V ^

V V V

V**-V

Leaf and reed remain

Mammalian bone

Silicified tree trunk

400 -

a— a-— a

■ Evaporitic facies

Reefal facies

Silty marly facies

Sandy facies

Gravelly facies

350

^4-

a—Q—tJH — £~

g_g.

■ o O, O.o'

O—P. o o. ‘&Q\

. 00 *o)

V V V

V v

4 !—*

Siliceous silt and clay

Basalt and tuff

Dogali

terraced alluvial gravel and cobble Hubbet

volcano-lacustnne succession

lower DOGALI

alluvial to shallow marine succession

_ upper DOGALI _

Gypsum

I !•! I Patch reef limestone

Fluvio-deltaic sand

1^ ol Alluvial gravel and cobble

|I~I| Coastal plain mud

High-density flow gravel

Siliceous lacustrine deposits

Basalt flow

Paleosoil HS Highstand systems tract

Bioturbation

TS Transgressive systems tract

Hummoky-cross stratification 7

Trouqh-cross stratification . . . .

a LLS Late lowstand systems tract

Fig. 7.— Summarized lithology of the Dogali Formation (a) and schematic cross section with the sequence stratigraphy

interpretation (b).

Fig. 7 .— Lithologie de la Formation Dogali (a) et coupe schematique avec Finterpretation en terme de stratigraphie

sequentielle (b).

Source: MNHN. Paris

JURASSIC AND NHOGENE SEDIMENTATION FROM ERITREA

207

volcanites. Scattered paleocurrent data inferred from imbricated or clustered pebbles and cross

laminations suggest main flows toward east;

— sandy facies. Cross or horizontally laminated and coarse to fine grained light yellow to green

sands in beds up to 2 m thick. Pebbles, clay clasts, mollusc shells and root moulds are locally frequent.

Burrows and pervasive bioturbation and subordinate wave ripples and hummocky laminations are

common features in these sands. Some low angle cuneiform laminations, enhanced by well-sorted and

imbricated flat clasts also occur;

— silty marly facies. Pale green massive silty marls occur as 0.5-1.5 m thick beds. They are

laminated, nodular and bioturbated and contain fine comminuted plant debris, shells of Melanopsis and

carbonate nodules (caliches). Abundant gypsum veins and nodules locally are also present. Red and

purple mottling, manganese spherules and caliche nodules mark pedogenized and rooted horizons;

— patch reef facies. Lenticular, 1.5-2.5 m thick, patch reef bodies occur at two levels in the

succession. Massive coral colonies are in growth position and float in a carbonate matrix. Stromatolitic

limestone fills the cavities within the corals. The reef covers calcilutite pale yellow limestone beds that

contain sparse pelecypods, single corals, cherty nodules, hematitic veins and dolomite crystals;

— evaporitic facies - Two levels of gypsum occur in the upper part of the measured section. The

lower one, 20 m thick, is made of nodular alabastrine gypsum capped by an alternance of laminated

microcrystalline gypsum and shales. This level is covered by coralline limestone and pebbly sandstone.

The upper one, 2 m thick, lays on pale green bioturbated marls capped by 0.5 m of dolomite-gypsum-

marl laminae. Gypsum shows nodular alabastrine and chicken-wire structure with ghosts of 5-10 cm

selenite "cavoli" crystals formed in hypersaline waters. On top enterolithic nodular structures and

vertical veins filled with crystalline gypsum occur. This gypsum horizon makes upward transition to red

clay and to crystalline gypsum alternating with silt.

Age of the Dogali Formation: K-Ar datings of basaltic rocks gave 24.0 to 14.0 Ma (KAZMIN, 1975),

but the location and stratigraphic position of samples is uncertain. More recently. DRURY et al. (1994)

report an Ar-Ar age of 28 Ma for the base of the Trap basalts below the Dogali Formation and an age of

18 Ma for a basalt flow in the lower Dogali Formation. A further indication of an early Miocene age for

the basal levels of this formation is given by the newly discovered Deinotherium bones found as clasts

in a pebbly channelized body. This finding recalls that of a Deinotherium tooth found in the lignite

levels interbedded in the Trap Basalts of the Eritrean plateau near Mendefera Adi Ugri (VlALLI, 1966).

Corals (Orhicella ellisiana , Cyphastrea corrugata, Plerastrea cf. profunda . Plerastrea saheliana,

Porites sp., Orhicella apenninica, Orhicella reussiana, Orhicella defrancei) and marine molluscs

(Pecten sp. and Ostrea sp.) suggest a middle to late Miocene age (ZUFFARDI-COMERCI, 1936 and

MONTANARO GALLITELLI, 1939) for the upper Dogali Formation.

The desset boulder beds

The Desset Boulder Beds form a continuous hilly chain extending from the Sudanese border to the

Zula Gulf (PORRO, 1936). They consist of very coarse pebbles deriving from basement. Mesozoic rocks

and Trap suite with outsized clasts up to 100 cm in diameter generally concentrated in the upper portion

of the beds. They are massive, poorly organized with abundant microgranule and coarse grained sandy

matrix. Beds are lenticular, amalgamated, structureless and rest above erosive irregular surface. They

form different terraces connected with distinct sedimentary cycles. Plio-Pleistocene marine fossils were

collected by BALDACC1 (1891) in the coastal outcrops of these deposits near Massawa.

Sequence stratigraphy analysis

In order to obtain information on the tectonic, environmental and climatic changes that have

controlled and guided the depositional processes, a sequence stratigraphy analysis has been carried out

on the Dogali Formation. The study has been limited to the upper part of this unit, where favourable

outcrops allow to recognize several unconformity-bounded sequences (Fig. 7). As above mentioned,

these are characterized by cyclic arrangement of facies associations that can be interpreted in terms of

sequence stratigraphy models.

At least four sequences (SI to S4; Fig. 7), 30 to 70 m thick, have been distinguished in the succession

208

MARIO SAGRI ET AL.

cropping out in the Desset River valley. Each sequence is formed of basal coarse-grained fluvial

deposits (facies a) which mark progradation and aggradation in the coastal area during the late lowstand

(LLS). These deposits pass gradually upward to sandy-muddy fluvio-deltaic (facies a, b) or coastal plain

sediments (facies b and c formed during the transgressive phase TS. The maximum marine ingression

and sea-level stillstand is recorded by patch reef limestone (facies d). The highstand deposits (HS) are

generally lacking and the typical fining and deepening upward (FU) trend is abruptly truncated by the

LSS deposits of the overlying sequence. Only in the lowermost sequence (SI) the presence of deltaic

deposits on top of the reefal limestone suggests a possible highstand. The symmetric depositional trend

evidenced by the deposits in SI could stem from an eustatic control on the base-level (i.e. sea-level)

fluctuations, while the FU trends in sequences S2 to S4 can be related to fast increase of accommodation

space in the basin due to an high rate of tectonic subsidence. A climatic signal (increasing aridity)

superimposed on a base-level fluctuation is evidenced by evaporitic deposits (facies e) of S4.

Structural remarks

The Dogali Formation dips homoclinally to the northeast around 10°-15° beneath the Desset Boulder

Beds with an incised irregular boundary. The occurrence of polymictic conglomerates in the Desset

Boulder Beds with clasts of basement. Mesozoic rocks and volcanics marks clearly the onset oi major

extensional faulting and uplift of the Eritrean plateau along the escarpment. This is connected with an

important tectonic pulse during Pliocene times that rejuvenated the relief causing deep incision by rivers

and deposition of Boulder Beds fanglomerate (BEYTH, 1978; DRURY el al ., 1994).

THE DANAKIL FORMATION IN THE NORTHERN AFAR DEPRESSION

According to GARLAND (1980). the Danakil Formation is an inhomogeneus unit varying in lithology

and in thickness from place to place. The lower to middle portion is formed by well consolidated

conglomerates, sandstones and mudstones intensely red to violet coloured, reflecting alteration under

very wet climatic conditions. Beds of basalt are commonly interbedded in the sedimentary section. The

unconformably overlying portion consists of sands, pebbles and clays with interbedded poorly cemented

limestones rich in marine gastropods, corals and pelecypods. Locally, the sediments are stuffed with

pillowed and spilitic basalt subaqueous lava flows, rhyolites and ignimbrites.

The Danakil Formation, with a maximum thickness of 1,000 m, rests on the basement and Mesozoic

rocks of the Danakil Horst and of the Eritrean escarpment.

The outcrops of the Danakil Formation that are close to the foot of the escarpements are overlain

with an erosional contact by multiterraced Boulder Beds fanglomerates coming from the shoulders of

the depression.

The danakil formation in the buia section

A thick succession composed of predominant grey to whitish silts and sands has been examined near

Buia at the northernmost apex of the Afar Depression and at the foot of the Eritrean escarpment (Fig. 8).

It outcrops in isolated patches beneath terraced younger sediments. Its lithology and geological setting

are reminiscent of the Neogene Danakil Formation (or Red Series) which outcrops widely further south

in the Afar Depression (see geological maps by BANNERT et al.. 1970; Barberi el al.. 1972; KAZMIN,

1975; MERLA et al.. 1979; GARLAND, 1980). Although a physical continuity has not been ascertained

nor any firm correlation has been attempted, we provisionally refer Buia section to the upper portion of

the Danakil Formation.

Composite sections have been investigated 5 km south of Buia in the Badda region near the Alut

water well. This 500 m thick section is possibly the upper portion of the Danakil Formation and consists

of predominantly terrigenous deposits laid down in lacustrine and fluvio-deltaic environments (Fig. 8).

Four main lithofacies have been recognized.

— Green, pale green siltstones and mudstones, generally massive and bioturbated in beds 0.5 to 1.5

m thick are distributed along the entire section. Current and wavy ripples and flaser bedding locally

occur. Thin paleosols horizons appear as pale brown, pink to mottled levels with nodular calcareous

concretions (caliches), and root marks. Bioturbated levels are characterized by abundant shell remains of

Source: MNHN. Paris

JURASSIC AND NEOGENE SEDIMENTATION FROM ERITREA

209

Fig. 8.— Lithology and environment interpretation of the upper Danakil Formation in the Buia area. Detailed logs of some

typical levels are added.

FlG. 8 .— Lithologie et environnement de la Formation Danakil superieure dans la region de Buia. Logs detailles de quelques

niveaux particuliers.

Source: MNHN. Paris

210

MARIO SAGRI ETAL.

Melanopsis , few vegetal organic matter and Fe concretions. These thinly laminated deposits indicate

very low energy environment with rhythmic settlement of clastic material.

— Beds of 2 to 4 m of grey to pale green and pale brown pebbly sands, massive with well-developed

trough-, cross- and plane-laminae and wavy ripples repeat along the entire section. Locally, the beds

show lenticular shapes and erosive boundaries. Normal reverse and symmetric grading are common

features together with well-rounded pebble clasts and clay chips scattered in the sandstone. Cherty

lenses, gastropods, rooted horizons and intense vertical burrowing are sometimes abundant, chiefly in

massive beds. Fluid escapes, dish structures and convolute laminations indicate intense clastic supply

and quick squeezing of interpore water.

Sands occur either as single beds or stacked in amalgamated lenticular and channelized bodies, 4 per

20 m in size, displaying erosive basal contact and flat top. Lenticular upward convex bodies, 1 to 3 m

thick, composed by upward thickening or symmetric sequences represent sandy lobes rooted and

intensely bioturbated.

— White silty calcareous marls, 0.2-1.1 m thick, occur in the upper portion of the section. They are

massive, nodular, bioturbated and contain scattered clay chips and Melanopsis shells.

An extensive key bed occurs interbedded in the marls. It consists of ostracod-bearing calcisiltites, 50

cm thick, with questionable ring-shaped, "donuts"-like structures. The latter are 15 cm in diameter, 5 cm

in cross sections and protrude into the underlying marly levels.

— Pale brown to yellowish medium-fine sands in beds 0.5-1.5 m thick. They are massive, graded

with clay chips and locally are organized according to the Bouma sequences.

According to our preliminary investigations, clear depositional sequences could not be defined.

However, facies associations allow to recognize deltaic and lacustrine environments (Fig. 8). At the

base, the prevailing of pebbly sands and silty facies indicate a delta plain environment with lenticular

bodies filling distributary channels. Up in the succession silts and clays bearing vegetal remains prevail

and indicate a swamp environment. Thick sandy bodies again testify delta plain deposits followed

upward by frankly central lacustrine facies characterized by marly limestones and turbidite beds. The

last 100m of the section consists of delta front and delta plain deposits representing a regressive trend.

The topmost beds are intensely bioturbated with large roots and burrows.

An angular unconformity of regional extension marks the contact between the Danakil Formation

and the overlying Boulder Beds.

As to the age of the Danakil Formation, radiometric datings come from the Ethiopian portion of the

Afar Depression. Intercalated basalts toward the base and the top of this unit gave early Miocene and

late Pliocene ages, respectively (25 Ma and 2,5 Ma; Bannert et al ., 1970; BARBERI et al ., 1972).

During our investigations five fossiliferous levels in the Buia section (Fig. 8) yielded an abundant

Vertebrate fauna with bones of elephants, hippos, antelopes, bovids, equids, suids, hyaenids and

crocodiles, and a Homo skull (see announcements MORELL, 1996). Although still under study, the

fossil association seems to indicate a wooded savannah environment and a early Pleistocene age for the

lacustrine and fluviodeltaic deposits. At that time the landscape was fundamentally different from the

present Danakil barren desert since the Eritrean plateau had not yet risen to its present height and was an

undulated smooth lowland. Researches in progress will hopefully contribute to the understanding of the

relations between climatic changes and mammals evolution (see MENOCAL, 1995).

Boulder beds

Boulder Beds similar to those found in the Dogali area rest on the Danakil Formation with an

erosional and unconformable surface (Fig. 8). They consist of boulders and cobbles of basement and

Mesozoic rocks and contain outsized boulder up to lm in diameter. They have the same climatic and

tectonic significance of those outcropping in the Dogali area. Similarly, they have originated from

catastrophic floods discharging materials derived from active fault scarps bordering the Eritrean Plateau.

The age of the Boulder Beds is Pleistocene.

Source: MNHN. Paris

JURASSIC AND NEOGENE SEDIMENTATION FROM ERITREA

21 1

Fig. 9.— Calcareous nannofossils in the Antalo Limestone of the Adeilo area.

The abbrevations CL and TL denote cross-polarized light and transmitted light, respectively; all magnifications x 2800.

a-b: Watznaueria barnesae (Black); a: distal view CL, Mahur section, b; TL as a. c-d: Watznaueria manivitae Bukry; c:

distal view CL, Assale section, d: TL as c. e-f: Watznaueria contracta (Bown & Cooper); e: distal view CL, Assale

section, f: TL as e. g-j: Cyclagelosphaera wiedmannii Reale & Monechi; g: distal view CL, Mahur section, h: TL as

figure g, i: distal view CL. Mahur section, j: TL as i. k-1: Podorhabdus grassei Noel; k: distal view CL. Mahur section.

1: TL as k. m-n: Discorhabdus ( Palaeopontosphaera ) dorsetensis Varol & Girgis; m: distal view CL, Mahur section, n:

TL as m. o-p: Stephanolithion bigotii maximum Medd; o: distal view CL, Mahur section, p: TL as o.

Fig. 9.— Nannofossiles des calcaires d'Antalo de la region d'Adeilo.

CL et TL = lumiere polarisee et lumiere transmise; toutes les photos sont x 2800. a-b : Watznaueria barnesae (Black);

a : vue distale CL. coupe de Mahur. b : TL idem a. c-d : Watznaueria manivitae Bukry; c : vue distale CL. coupe

d'Assale. d : TL idem c. e-f : Watznaueria contracta (Bown & Cooper); e : vue distale CL coupe d Assale. f: TL idem

e- g j •' Cyclagelosphaera wiedmannii Reale & Monechi; g: vue distale CL. coupe de Mahur. h : TL comme g. i: vue

distale CL. coupe de Mahur. j : TL idem i. k-l: Podorhabdus grassei Noel; k : vue distale CL. coupe de Mahur, I: TL

idem k. m-n : Discorhabdus (Palaeopontosphaera) dorsetensis Varol & Girgis ; m : vue distale CL, coupe de Mahur, n :

TL idem m. o-p : Stephanolithion bigotii maximum Medd; o : vue distale CL. coupe de Mahur, p : TL idem o.

Source: MNHN. Paris

212

MARIO SAGRI ETAL.

CONCLUSION

The main results obtained from our investigations in the Danakil horst, Sahel lowland and Danakil

(Afar) Depression can be summarized as follows:

— stratigraphical and sedimentological definition of six units in the Antalo Limestone and

paleoenvironmental recognition of repeated cycles within a ramp system;

— significant biostratigraphical data, mainly based on ammonites, calcareous nannoplankton and

foraminifers that assign a late Callovian to early Kimmeridgian age to the Antalo Limestone;

— recognition of a deltaic episode (the Simi Koma key bed) within the marine Antalo Limestone;

— interpretation according to sequence stratigraphy of the Antalo Limestone and recognition of

seven depositional sequences, systems tracts and significant surfaces;

— geological map of the Adeilo area showing the areal extension of the different units in the Antalo

Limestone and how intensely they have been deformed (also under a transcurrent regime) during the

Neogene development of the Southern Red Sea and Afar rift systems;

— basin and facies analyses in the Dogali and Danakil formations;

_recognition in the Dogali-Desset section of a lower (siliceous lacustrine deposits and basalts) and

an upper (fluviodeltaic to shallow marine) portion of the Dogali Formation.

_facies analysis evidences four depositional sequences mainly connected with relative sea-level

fluctuations;

— new paleontologic record ( Deinotherium) for an early Miocene age of the Dogali Formation;

_facies analysis in the upper portion of the Danakil Formation with the assessment of fluviodeltaic

and lacustrine environments;

— findings of an abundant early to middle Pleistocene vertebrate fauna and, in particular, of a Homo

skull, that aTso provide paleoclimatic and paleoecologic information of remarkable significance for

discovering the origin and interpreting the evolution of the human species.

ACKNOWLEDGEMENTS

All the authors participated into the 1994 and 1995 field campaigns to Eritrea and jointly contributed

to data interpretation and presentation. In particular, M. SAGRI, E. ABBATE, M.L. BALESTRIERI, M.

Benvenuti, P. BRUNI, M. Fazzuoli, M. Papini and Tewelde Medhin TECLE: regional geology,

stratigraphy and sedimentology; A. AZZAROLI, G. FlCCARELLl and L. ROOK: vertebrate paleontology;

G. Pavia: ammonites; M. MARCUCCI: foraminifers; V. REALE: calcareous nannofossils.

The authors thank the Peri-Tethys Program, the Italian National Research Council, the European

Commission, the Italian Ministry for the University and Scientific and Technological Resaerches and the

University of Florence for supporting their researches. The Department of Mines of the Eritrean

Ministry of Energy, Mines and Water Resources and the Italian Embassy in Asmara are to be

acknowledged for providing field facilities and logistic assistance. Thanks are also due to two

anonymous reviewers that substantially improved an early version of this paper. Professor Michel

SEPFONTAINE kindly helped in benthic foraminifers determinations. Paolo GHETTI and Alessandro

ABBATE provided a generous assistance in the field.

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Tectonic evolution of the Southern Tethyan

margin in Southern Tunisia

Samir BOUAZIZ" 1 , Eric BARRIER 121 , Jacques ANGELIER' 2 ',

Pierre TRICART 131 & Mohammed M. TURKI 141

<h Departement dc Geologie, Ecole Nationale d’Ingenieurs de Sfax, B.P. W, 3038 Sfax, Tunisie

<2) Departement de G6otectonique, URA CNRS 1759, Universite P. et M. Curie, 75252 Paris Cedex 05, France

(3) Institut Dolomieu. Geologie et Mineralogie, UPRES CNRS 5025, Rue M. Gignoux, 38031 Grenoble, France

,4 ' Universite de Tunis II, Faculte des Sciences de Tunis, Departement de Geologie, 1060 Tunis, Tunisie

ABSTRACT

Southern Tunisia is located at the northern edge of the Saharan platform and South of the folded Atlasic domain. It is

characterized by major sedimentary basins providing excellent outcrop conditions. The stratigraphic section consists of an

almost complete sequence that ranges in age from Late Permian to Quaternary where fractures exist in all geological

formations. Two distinct domains have been investigated: the stable platform and the Southern pre-Atlasic domain. An

extensive analysis has been carried out in 354 sites. It allows the reconstruction of the tectonic evolution of the Southern

Tunisian domain since Late Permian in term of paleostress tensors (434 have been computed). From Permian to Present, the

tectonic evolution of Southern Tunisia includes the following succession of tectonic events: a) from Late Permian to Triassic,

during the break up of Pangea, extensional tectonics dominated; b) from late Camian to Early Cretaceous, the tectonics were

characterized by sub-meridian extensions related to the Africa-Eurasia divergence; c) in Late Cretaceous and Paleogene

tectonics remained distensive, but the directions of extension rotated several times; d) the Neogene time is marked by

compressional events resulting from the Africa-Eurasia convergence. The major event is the NW-SE Atlasic compression,

mainly Tortonian in age. During Plio-Quatemary, 3 other minor compressional events, trending 020-040, 080-090 and 100-110

affected the studied area.

RESUME

Evolution tectonique de la marge sud-tethysienne dans le Sud-Tunisien.

La Tunisie meridionale est situ£e sur la bordure nord de la plate-forme saharienne, au Sud du domaine plisse atlasique.

Cette region est caracterisee par la presence de bassins sedimentaires fournissant d'excellentes conditions d'affleurement. La

serie stratigraphique est presque continue du Permien superieur au Quatemaire et presente une fracturation dans tous les

niveaux. Deux domaines distincts ont ete etudies : la plate-forme stable et le domaine pre-atlasique meridional. Une analyse

detaillee de 354 sites a permis de reconstituer revolution tectonique de la Tunisie meridionale depuis le Permien superieur en

terme de tenseurs des paleo-contraintes (434 ont ete calcules). Du Permien a 1'Actuel 1'evolution de la Tunisie meridionale

presente la succession suivante d’evenemcnts tectoniques : a) du Permien superieur au Trias, pendant l'eclatement de la Pangee

les tectoniques distensives dominent; b) du Carnien superieur au Cretace inferieur, la tectonique etait caracterisee par une

extension sub-meridienne liee a la divergence Afrique-Eurasie ; c) au Cretace superieur et au Paleogene, le contexte tectonique

Bouaziz, S., Barrier. E., Angelier, J., Tricart, P. & Turki, M.M.. 1998.—Tectonic evolution of the Southern Tethyan

margin in Southern Tunisia. In: S. Crasquin-Soleau & E. Barrier (eds), Peri-Tethys Memoir 3: stratigraphy and evolution of

Peri-Tethyan platforms. Mem. Mus. natn. Hist, nat., 177 : 215-236. Paris ISBN : 2-85653-512-7.

Source: MNHN. Paris

216

SAMIR BOUAZIZ ETAL.

reste distensif mais les directions d'extension tournent plusieurs fois ; d) le Neogene est marque par des evenements

compressifs resultant de la convergence Afrique-Eurasie. Levdnement majeur est la compression NW-SE. due phase

atlasique", d'age tortonien. Au Plio-Quaternaire. trois autres phases mineures de compression, respeetivement orientees 020-

040, 080-090 et 100-110, ont ete mises en Evidence.

INTRODUCTION

In Southern Tunisia, the Saharan platform constitutes the northern part of the stable African plate

(Fig. 1). During Mesozoic and Cenozoic times, this domain formed the southern Tethyan margin at the

boundary between Western and Eastern Tethys. Because of this particular location, the tectonic

evolution of the northern African margin in southern Tunisia has been successively governed since Late

Permian by the break up of Pangea (AUBOUIN et al., 1980), the opening of the East Mediterranean basin

(BIJU-DUVAL et al., 1976; DERCOURT et a!., 1993; RlCOU, 1994), and the convergence between Eurasia

and Africa (BOUSQUET, 1976; LETOUZEY & TREMOLIERES. 1980; PHILIP. 1987). From this long tectonic

evolution results the present structural complexity of Southern Tunisia.

The tectonic evolution of the African margin in Tunisia is governed by inherited structures

(DURAND-DELGA, 1981; BOU1LLIN, 1986: TRICART et al., 1994) reactivated during the Neogene and

Quaternary compressional events. The northern part of the African plate is constituted from South to

North by several domains (Fig. 1): the stable Saharan platform, the Atlasic folded domain, the Tellian

nappes and the African basement only known in Kabylia. Southern Tunisia is a transition zone between

the tabular domain and the South folded Atlasic belt.

In this paper, we present a reconstruction of the paleostress evolution from Late Permian to Present

based on the analysis of brittle deformations. Paleostress analysis of brittle structures is a powerful tool

to reconstruct the tectonic evolution of sedimentary basins. The investigated area is Southern Tunisia

Fig. 1.— Location map and main tectonic features of the northern african margin. 1, African basement; 2, Tellian nappes; 3,

Saharan Platform; 4, Late Miocene basins; 5, Kabylian front; 6, Tellian front; 7. Saharan flexure; 8, Folds; 9. Faults

Box; location of Figure 2.

FlG.I — Localisation de la zone etudiee et principals structures techniques. I, socle ; 2. nappes telliennes ; 3, plate-forme

saharienne ; bassins Miocene superieur ; 5, front des nappes kabiles ; 6, front des nappes telliennes ; 7, flexure sud-

saharienne ; 8, axes de plis ; 9. principals failles. Encart: localisation de la figure 2.

Source: MNHN. Paris

TECTONIC EVOLUTION OF SOUTHERN TUNISIA

217

(Figs 1,2) including (a) the stable Saharan platform, located South of the Chott Range and of the Gabes

Gulf, and (b) the folded southern pre-Atlasic domain of the Chott and Gafsa ranges. The tectonic

framework of these domains has been discussed by many authors in both Algeria (LAFFITE, 1939;

CASTANY, 1954; AlSSAOUl, 1984; VlALLY et al., 1994) and Tunisia (ZARGOUNI, 1985; BOUKADI,

1994). Different interpretations are made on the significance and age of major structures. A new

calendar of tectonic evolution is proposed where tectonic events are considered in term of stress. Several

periods are considered following the main stress regimes and the geodynamical context. We precise the

complex tectonic history resulting from this evolution. For each period, the results are discussed in the

context of the geodynamical evolution of the Tethyan Ocean.

GEOLOGICAL SETTING: THE SOUTHERN TETHYAN MARGIN IN SOUTHERN TUNISIA

In Southern Tunisia, the North African margin contains two distinct domains: (a) the Dahar Plateau;

a platform constituting the northern border of the stable Saharan platform, and (b) the folded southern

Atlasic domain, often called “pre-Atlasic domain ”, and constituting the deformed foreland of the

Tellian Atlasic chain (Figs 1, 2). The synthetic stratigraphic section of these two domains consist of an

almost complete sequence that ranges in age from Late Permian to Quaternary. In both domains, outcrop

conditions are excellent, providing numerous good sites for brittle tectonic analysis.

The stable platform: the Dahar plateau

The Dahar Plateau is a gentle large scale monocline dipping southward of l°-2° that may be

considered as tabular at regional scale. The Dahar, belonging to the Saharan platform domain, is

constituted by Late Paleozoic to Mesozoic sequences ranging in age from Late Permian to Late

Cretaceous (Fig. 2a). These sequences constitute the southern Tunisian basin of BUSSON (1967). They

are well exposed along continuous cliffs bordering the plateau to the North along the Jeffara coastal

plain, whereas they are covered to the West and to the South by the dunes of the Eastern Saharan Erg.

The sequences consist mainly of carbonates where dolomitic facies dominate. However sandstone

layers are also common. Three periods of sedimentation can be differentiated separated by two major

unconformities, Carnian and Albian in age, related to regional tectonic events:

— the Late Permian to Late Triassic cycle is characterized by a high subsidence rate during Late

Permian (BUSSON, 1969), then diminishing in Early-Middle Triassic when started the continental

sedimentation marked by the deposition of sandstones and clays (BUSSON, 1967; BOUAZIZ et al., 1987;

MOCK et al., 1987; KAMMOUN et a !., 1994). In the Triassic sequences, syn-depositional tectonics were

responsible for local angular unconformities, abrupt thickness variations and facies changes that are

frequently observed (BOUAZIZ, 1995). They affected the Early-Middle Triassic sandstones facies as well

as the Late Triassic carbonate facies;

— the marine early Carnian transgression, which began the second cycle, covered the whole platform

in Southern Tunisia and Western Libya above the Sidi Stout unconformity (BUSSON, 1967; MELLO &

BOUAZIZ, 1987). The marine environment lasted during the entire Jurassic and Neocomian. Extensional

faulting caused thick sedimentary accumulations in E-W trending basins (Tataouine and Chott basins).

The Jurassic extension induced high rate subsidence in which thick evaporites were deposited during the

Liassic. Sedimentation is contrasted between plateaus and basins (BOUAZIZ, 1986; BEN ISMAIL et a !.,

1989). In the first, carbonate facies dominate, whereas the second are characterized by a deeper

environment. In Early Cretaceous a withdrawal of the marine domain began but marine conditions still

prevailed;

— the second marine transgression that started in the Albian time (Vraconian). It expanded quickly

during the Cenomanian before reaching its maximum in the Turonian time. Such transgression is also

known in central Tunisia, in Algeria and in Libya (BUROLLET & BUSSON, 1983). The Late Cretaceous

deposits are constituted of a 200 m to 300 m thick carbonate sequence that significantly increase in the

Jeffara Coastal plain. Evidences of synsedimentary deformations (slumps, fault breccias, thickness and

facies variations, local unconformities) are common within the Late Cretaceous series, particularly in the

Cenomanian and in the Coniacian-Santonian deposits.

218

SAMIR BOUAZIZ ETAL.

Source: MNHN, Paris

TECTONIC EVOLUTION OF SOUTHERN TUNISIA

219

The folded domain: the chott and gafsa ranges

This domain constitutes the folded foreland south of the Alpine chain (Fig. 1). In this area, far south

ot the allochtonous Tellian domain, the fold geometry is characterized by a succession of tight anticlines

separated by large box synclines (Fig. 2a). The anticlines are large ramp-related folds of various types;

i.e. fault-bend folds and fault-propagation folds (OUATTANI et al. y 1995). This folding is generally

considered as resulting from tectonic inversions of inherited extensional structures during the Cenozoic

compressional events (ZARGOUNI, 1985; BOUAZIZ, 1995). The pattern of previous faults probably

causes the complex fold geometry observed in the Chott and Gafsa Ranges (Fig. 2a). In this region, E-W

trending elongated folds predominate (Kebili Tebaga Djebel and El Asker Djebel in the Chott Range;

Alima Djebel and Chemsi Djebel in the Gafsa Range). NE-SW to ENE-WSW trending folds are also

common, but shorter (10 km to 30 km long) than the major E-W ones (60 km to 130 km long). In this

area, there is no evidence for two superimposed fold-thrust events as described in Central Tunisia

(SOYER & TRICART, 1989). In addition, NW-SE trending lineament, including the right lateral Gafsa

fault system (ZARGOUNI, 1985), cuts the region from the Algerian border to the Gabes Gulf (Figs 1, 2).

This complex domain is a key for understanding the Cenozoic tectonic evolution and complete the

information obtained from the investigation of the Mesozoic sequences of the stable platform.

The South Atlasic domain in the studied area is characterized by a thick Cenozoic sequence, missing

in the stable platform. These Cenozoic deposits conformably overly the Late Cretaceous sequence

(M’Rabet, 1981), mainly represented by carbonate facies ranging in age from Turonian to

Maastrichtian. The most common Mesozoic facies is the Campano-Maastrichtian “ Abiod Formation ”

(BUROLLET, 1956) forming the cores of the anticlines. The underlying Jurassic sequences, known from

subsurface data, consist in epicontinental deposits (marls and marly-sandstones) more than 2000 m thick

in the Chott basin where the Jurassic subsidence is maximum. Two-distinct periods of sedimentation can

be differentiated during the Cenozoic: Paleogene sedimentation remains marine, whereas during the

Neogene continental basins developed. The Paleocene-Eocene sequence begins by Paleocene clays

underlying the Ypresian Metlaoui formation represented by Carbonates and phosphate layers (SASSI,

1964). The Metlaoui formation reaches one hundred meters in the center of the basin. This sequence

ends by late Eocene gypsum. In the studied area Oligocene is missing. The continental Mio-Pliocene

molassic series rest unconformably over the pre-Neogene formations. They are characterized by a

detritic sedimentation including sands, sandstone beds, conglomerates and clays. They include the

Sehib, Beglia and Segui formations, respectively Aquitanian, Serravallian-Tortonian and Plio-

Quaternary (BlELY et al ., 1972; MALLOULI et al ., 1987). The thickness of these series reachs several

hundred of meters in the synclines where they are well preserved. The most recent formations are the

sporadic “ Villafranchian ” (late Pliocene-early Pleistocene) crust (VAUFREY, 1932; COQUE, 1962;

BEN OUEZDOU, 1994) and Late Quaternary river terrasses.

FlG. 2.— A: geological map of Southern Tunisia.

a. Dahar plateau - Stable Platform; b: South Atlasic belt; 1, Permian; 2. Early-Middle Triassic; 3. Camian; 4. Camian-

Norian; 5. Rhetian; 6, Rhetian-Early Liassic; 7. Pliensbachian; 8. late Liassic; 9. Bathonian-Bajocian; 10. Callovian; 11.

Malm-Neocomian; 12. Barremian; 13. Aptian-Albian; 14. Cenomanian-Turonian; 15, Coniacian-Santonian; 16,

Campanian-Maastrichtian; 17, Paleocene-Eocene; 18. Mio-Pliocene; 19, Quaternary.

B: Location of the sites of fault measurements. AST, South Tunisian Fault.

FlG. 2.— A : carte geologique de la Tunisie meridionale.

a, Plateau du Dahar, plate-forme stable ; b, chaine sud-atlasique ; l, Permien ; 2, Trias inferieur et moyen : 3,

Carnien ; 4, Carnien-Norien ; 5, Rhetien ; 6. Rhetien-Lias inferieur ; 7, Pliensbachien ; 8. Lias superieur ; 9,

Bathonien-Bajocien ; 10, Callovien : 11, Malm-Neocomien ; 12, Barremien ; 13, Aptien-Albien : 14, Cenomanien-

Turonien ; 15, Coniacien-Santonien : 16, Campanien-Maastrichtien ; 17, Paleocene-Eocene ; 18, Mio-Pliocene ; 19,

Quote rnaire.

B: Localisation des sites etudies. AST, faille sud-tunisienne.

220

SAMIR BOUAZIZ ETAL.

BRITTLE TECTONIC ANALYSIS

An extensive brittle tectonic analysis has been carried out in a total of 354 sites of the South Tunisian

domain in order to reconstruct the stress pattern evolution since the end of Paleozoic. Location of sites

are shown on figure 2b. Brittle tectonic analysis includes (a) analysis of fault-slip data sets, associated

with paleostress tensor reconstructions, and (b) study of joint populations, more common than fault

populations in the stable platform. These sites are distributed in all formations ranging in age from Late

Permian to Pleistocene, except in the Late Triassic-Liassic gypsum. Site are located on the stable

platform (72% of the sites), and on the folded Atlasic domain (28%). Some formations have been more

extensively studied because either they are widespread in both domains, like the Late Cretaceous

sequences (163 sites), or they are of particular interest like the Triassic deposits of the platform (41 sites)

or the Paleogene sequences of the Gafsa basin (31 sites). The Jurassic formations of the platform, where

joint populations are common, have been also studied in detail (72 sites), as well as the Miocene

continental deposits (28 sites) that recorded the Cenozoic compressive events. The age of tectonic events

was established based (a) on the age of rock formation affected, and (b) on the succession of tectonic

features where deformation is polyphased.

Each site consists of several tens of tectonic features measured in a small area. They include mainly

faults with slickenside lineations, but also bedding planes, tension gashes and fold axes. The method

enables us to reconstruct the paleostress axes (maximum stress ol, intermediate o2 and minimum o3)

and the ratio 0=o2-o3 / ol-o3, between principal stress magnitudes, knowing the orientations and

senses of slip on faults acting during the same tectonic event. The principles, conditions of application

and limits of the methods have been described in detail by ANGELIER (1984, 1989, 1990). Many fault

sets are practically monophase and reveal simple conjugate patterns of neo-formed faults. However,

numerous sites, mainly located in the folded domain, display more complicated fault patterns including

oblique and inherited fault slips or polyphase faulting. Particular attention was focused on these

polyphase and superposed deformations resulting from different tectonic events, and on syndepositional

faulting providing direct stratigraphic dating. Many sites are polyphased including fault populations

resulting from distinct tectonic events. In the platform, where deformation rates are low, neoformed

normal fault systems and joint sets are often reactivated during later compressive events as strike-slip

faults or reverse faults and ramps. In the folded domain, where fault tectonic is more complex, we pay

particular attention to distinguish pre and post-tilting (or folding) fault populations in order to

reconstruct the chronology of faulting (and the paleostress evolution) with respect to the major

deformation phases. In the whole studied area, 434 paleostress tensors have been determined in 247 sites

(165 in the platform and 82 in the folded domain). The complete results of paleostress tensor

determinations have been reported in BOUAZIZ (1995). Two types of paleostress tensors were

distinguished: 44% of them correspond to strike-slip and compressional faulting (ol horizontal), and

56% to extensional faulting (c3 horizontal and ol vertical).

Joints have been extensively analysed in terms of nature and orientation. 107 sites of joints have been

analysed (92 in the platform and 15 in the folded domain). All type of joints have been observed

(tension, hybrid and shear types). The most common are tension joints, that we correlated often with

surrounding normal fault populations, or more occasionally with hybrid joints. In many case, we might

recognize the mechanical significance of the different sets of joints and thus identify at least one of the

three axis of paleostress. Relationships with fault populations show that most joint patterns are of

tectonic significance. In this case, they are reliable indicators of paleostress orientation.

TECTONIC EVOLUTION OF THE NORTH AFRICAN MARGIN IN TUNISIA

The tectonic history of the Saharan platform in Southern Tunisia since Late Paleozoic is controlled

by the regional geodynamical evolution characterized by 3 major periods: (a) the break up of Pangea

during the Permo-Triassic; (b) the divergence between Africa and Laurasia, resulting from the opening

of Tethys Ocean during the Late Triassic, Jurassic and Cretaceous times, and (c) the convergence

Source: MNHN. Pans

TECTONIC EVOLUTION OF SOUTHERN TUNISIA

221

between Africa and Eurasia, starting at the end of Cretaceous and developing during the all Cenozoic,

involving the closure of the Tethyan domain. Generally speaking, at the scale of the northern

Gondwanian margin, then of the northern African margin, the tectonic evolution is marked by the

inversions of basins mainly resulting from Mesozoic extensions.

The break up of pangea: the Permo-Triassic rifting

AND THE TRANSCURRENT MOVEMENTS

The Late Permian rieting

During the early stage of break up of Pangea, extensional stress fields affected wide areas of Laurasia

and Gondwana. They originated rift systems associated with high subsidence rates. In southern Tunisia,

the Late Permian reef carbonates crops out in the Medenine Tebaga (Figs 2, 3) where 850 m of sequence

is exposed (CHAOUACHI et al., 1987; M'RABET el al ., 1987; Razgallah el al., 1989). They constitute

the upper part of the 6000 meters thick sequence of the Jeffara basin (BAIRD, 1967; BUROLLET et al.,

1978) known from subsurface data (GlJNTZBOECKEL & RABATE, 1964). Northward the facies changes

from reef carbonates to more pelagic shaly deep sea deposits, whereas southward shallow water facies

and alluvial deposits mark the southern boundary of the basin. Subsurface data show that the Late

Permian Jeffara basin trends E-W to WNW-ESE. It ends abruptly westward where no Permian marine

deposits have been reported. Its eastern extension in Libya and offshore is unknown. The reef carbonates

of the bioherm complexes of the Medenine Tebaga are not a favorable facies for brittle tectonic analysis.

Nevertheless, several sites have been investigated in interbedded sandstone layers (Fig. 3). Only one site

of the Medenine Tebaga exhibits synsedimentary normal faults attesting for a Late Permian distension

(Fig. 3, diagram 2). Because the poor quality of the fault population, it is not possible to determine

accurately the Late Permian stress tensor. Joints are characterized by two sub-perpendicular major sets

of pre-tilting joints trending roughly N-S and E-W (Fig. 3, diagrams a, 1 and 3). Minor set of NW-SE

trending joints are also present. The geometry of joint sets pleads for tension joints resulting from a sub¬

meridian Late Permian extension. This assumption is supported by the same E-W general trend of the

basin. Minor directions may be due to later jointing. These evidences document a fault-bounded

subsiding basin, despite the major fault cannot be observed. In Arabia, the same age Palmyra rift is also

characterized by a thick clastic sequence, Carboniferous to Triassic in age. RlCOU (1994) suggests that

the Permian extension corresponds to an aborted rift that precluded the East Mediterranean basin.

The Early to Late Triassic extensions

We investigated the Early to Middle Triassic continental sequence, widespread in the cliff bordering

the Dahar Plateau to the north (Figs 2a, 3). It is a 800 m thick sequence, mainly composed of sandstones,

conglomerates and clays where crossbeddings and local unconformities are frequent (BUSSON, 1967;

BOUAZIZ, 1986, 1995). A detailed analysis of joint and fault populations indicates that a extensional

tectonic context predominated during this period. It is characterized by conjugate systems of

syndepositional normal faults where slickenside lineations are rare due to the granulometry of the

formations, and to the synsedimentary character of the normal faults. Only few sites allowed to

determine paleostress tensor (Fig. 3, diagram 9). They indicate NW-SE to NNW-SSE extensions. In

Djebel Rehach, ENE-WSW to NE-SW trending cartographic normal faults, associated with

synsedimentary features are sealed by the early Carnian Mekraneb dolomites. Fractures are more

common than faults in the continental Triassic facies. The study of joint populations corroborates the

fault tectonic analysis. Two perpendicular sets of pre-tilting tension joints, trending NE-SW and NW-

SE, have been measured, in relation with NW-SE and NE-SW extension (Fig. 3, diagrams b, c and 4 to

6). NE-SW trending normal fault populations and joint sets dominate. That may indicate that the NW-

SE extension was the major tectonic event during this period, whereas the NE-SW one might be related

to inversions between the main stress axes g 2 and o3.

This Early-Late Triassic extension tectonic constitutes the last stage of the Permo-Triassic

extensional period. In the Mediterranean area, this epoch corresponds to a rifting period. Triassic crustal

extension is documented in the Mediterranean Tethys by the formation of deep water Triassic basins

222

SAMIR BOUAZIZ ETAL.

Fig. 3.— Geological map of the Jebel Tebaga of Medenine and main sites of fracture analyses.

1. Late Permian; 2, Early-Middle Triassic; 3, Late Triassic-Liassic; 4, Bathonian pp.-Oxfordian; 5. Barremian; 6,

Aptian-Vraconian; 7, Cenomanian; 8. Turonian; 9. Neogene-Quatemary; 10, Faults.

Schematic geological cross-section. 1. Late Permian; 2, Undifferenciated Permian; 3, Bathonian pp.; 4. Bathonian pp.-

Oxfordian; 5, Barremian; 6. Aptian; 7, Vraconian; 8. Quaternary.

Rose diagrams; a. in Late Permian; b. in Early-Middle Triassic; c, in Late Triassic; n=Number of joints. Fault diagrams:

Schmidt’s projection, lower hemisphere, fault planes as continuous lines, and slickenside lineations as dots with arrows

(convergent for reverse slips, divergent for normal slip, double for strike-slip), bedding planes as dashed lines, main

axes of paleostress as stars (5, 4 and 3-branches for Ol. O2 and 03), main directions of extension and compression as

large black arrows. Upper left comer: sites in the Late Permian sequence, other diagrams in Triassic formations.

Fig. 3 — Carte geologique du Jebel Tebaga de Medenine et localisation des principaux sites analyses.

1, Permien superieur ;2, Trias moyen-inferieur ; 3. Trias superieur-Lias ; 4. Bathonien pp.-Oxfordien ; 5. Barremien ;

6, Aptien-Vraconien ; 7, Cenomanien ; 8, Turonien ; 9. Neogene-Quaternaire ; 10, failles.

Coupes geologiques schematiques. I. Permien superieur; 2, Permien indifferencie ; 3, Bathonien pp. ; 4 , Bathonien

pp.-Oxfordien ; 5, Barremien ; 6, Aptien ; 7, Vraconien ; 8. Quaternaire.

Diagrammes de frequence des directions de joints ; a, dans le Permien superieur ; b. dans le Trias inferieur-moyen ; c,

dans le Trias superieur ; n=nombre de joints. Diagrammes de failles : projections de Schmidt hemisphere inferieur ;

lignes continues : plans de failles ; points avec des fleches: stries de glissement (convergentes : failles inverses,

divergentes : failles normales, doubles fleches : decrochements); tiretes : plans de pendage ; axes principaux des

paleocontraintes : etoiles (5 , 4 et 3-branches pour Oi, 02 and (73), fleches noires : directions principals d'extension et

de compression. Encart en haut a gauche : sites permiens (les autres diagrammes sont dans les formations triassiques).

Source: MNHN. Paris

TECTONIC EVOLUTION OF SOUTHERN TUNISIA

223

which formed at the expenses of former epicontinental platforms: in Algeria, Triassic extension induced

synsedimentary normal faulting bounding zones of high rate subsidence in which thick saliferous series

were deposited (BUSSON, 1970; BOUDJEMA, 1987); in Libya, a Middle-Late Triassic rifting produced

horsts and grabens, and a crustal thinning in the Ionian Sea (DEL BEN & FlNETTl, 1980). These

extensions are generally associated to the Permo-Triassic rifting of Western Pangea (RlCOU et al., 1986;

STAMPFLI et al., 1991).

The late Carnian transcurrent movements

The N-S regional Sidi Stout unconformity marks a major Late Triassic tectonic event. The late

Carnian Rehach dolomites rest unconformably over the whole Late Permian to Middle Carnian

sequence. The transgressive sequence overlies unconformably the Late Permian in the area of Djebel

Tebaga of Medenine, the only large scale structure of the Dahar (MATHIEU, 1949; BUSSON, 1967;

BouAZIZ, 1986), where the unconformity reachs 30°. Southward, the angle of the unconformity

appreciably decreases and the Rehach dolomites are almost concordant with the middle Carnian

deposits. This tilting is related to a folding in relationship with a shear zone, roughly trending E-W, that

can be locally observed in the vicinity of the Djebel Tebaga. Study of pre-tilting fault populations in

Late Permian to late Carnian formations shows a N 150°E trending compression recorded in the Early

Triassic sandstone layers (Fig. 3, diagram 7). In same site, two pre-tilting joint sets (Fig. 3, diagram 8),

trending N-S and 130-140, are probably conjugate sets of shear joints, linked to the same compression.

This Carnian folding is related to a large E-W right lateral strike-slip fault zone, known from seismic

profiles, that extends eastward offshore. It may be also connected westward with the long fault zone

bordering the Saharan platform. This shear zone may be part of the dextral transcurrent faults that

separated Gondwana and Laurasia which led north-western Africa to shift away from central Europe. In

Maghreb, this fault zone is also documented from field data in the Tizi-n-Test in Morocco (Mattauer

et al ., 1977) and is continuous on the other side of Atlantic (RlCOU, 1992).

The Africa-Eurasia divergence: the mesozoic extensions

In the stable platform in southern Tunisia, from late Carnian to Late Cretaceous only extensional

tectonics have been evidenced. From a geodynamical point of view, this period corresponds to the

opening phases of Tethys and Central Atlantic and to the divergence between Gondwana and Laurasia,

and then between Africa and Eurasia. Two main tectonic periods are distinguished, separated by a low

angle regional unconformity, early Albian in age.

The late carnian to early aptian n-s extension

According to our data, the deformations were exclusively extensional between late Carnian and early

Aptian. A submeridian extension dominated in the whole platform during this period. In Jurassic time,

the sedimentation was clearly controlled by 090 to 110 trending normal faults. Two main zones of

subsidence, the Foum Tataouine and Chott basins, separated by ridges, were bounded by major normal

faults. The consequences are the important variations in facies and thickness in the Jurassic formations,

especially during Liassic time. Figure 4 shows the main Jurassic normal faults of the Tataouine basins

sealed by the post-Aptian sequences. Diagrams illustrate the 090 to 110 trending conjugate normal fault

systems related to this extension. In the Jurassic, these fault systems are common and often associated

with synsedimentary features. Joint pattern illustrates also the Jurassic extension as shown in BARRIER

et al. (1993), BOUAZIZ^r al. (1994) and BOUAZIZ (1995). More than 40 sites of joints have been studied

in the Jurassic layers of the stable platform. Part of this population is composed of two sub¬

perpendicular sets: a main set, roughly trending 110, frequently formed first, and a minor set trending

020. Most of this joints are sub-vertical tension joints. In summary, between late Carnian and Aptian, the

tectonic is controlled by a major 000-020 extensional trend expressed by normal fault populations and

joint sets. From South to North, the tectonic instability increases from the stable platform to the

subsident troughs of the Chott basin.

In this period an extensional context have been described in Algeria (ELMI, 1977; BUREAU, 1984;

VlALLY et al., 1994) and in Central Tunisia during Early Cretaceous (SOYER & TRICART, 1989). In

224

SAMIR BOUAZIZ F.TAL.

FlG. 4.— Late Triassic to Lower Cretaceous sub-meridian extension in Southern Tunisia. 1, Late Permian-Triassic; 2, Liassic;

3. Dogger; 4. Malm-Early Cretaceous; 5, Vraconian; 6, Late Cretaceous; 7. Faults. Diagrams, same than in figure 3.

FlG. 4 .— Extension sub-meridienne Trias superieur-Cretace inferieur en Tunisia meridionale. I, Permien superieur-Trias ; 2.

Lias ; 3, Dogger; 4 . Malm-Cretace inferieur ; 5, Vraconien ; 6, Cretace superieur ; 7, failles. Diagrammes, me me

legende que figure 3.

Central Tunisia, the Jurassic basins are characterized by platform deposits controlled by normal faults

(TURKI, 1985; TOUATI, 1985; BEN AYED, 1986; SOUSSI, 1990). In the Tunisian Sahel, ELLOUZ (1984)

and BEN AYED (1986) mentioned the same extensional context associated with lava flows. In Libya, an

important rifting phase occured in Middle Jurassic time accompanied by magmatic activity (DEL BEN &

Finetti, 1980). In Morocco, a syntectonic sedimentation has been described (Robillard, 1979;

SEUFERT, 1988). Normal faults trending NE-SW, active during Jurassic, have been described in Atlas

(MATTAUER el al ., 1977; Laville, 1981). They have been related to the opening of the central Atlantic

responsible for the Maghreb transform zone along which strike-slip displacements and transtensional

opening occurred. An extension tectonic event affected the entire present North African margin. The

regional sub-meridian extensional context, lasting from late Carnian to Aptian, has been considered by

several authors as resulting from a rifting (BUREAU, 1984; TURKI, 1985; SOYER & TRICART, 1989;

ALOUANI et al ., 1992). More than a simple rifting, we may consider this general extension along the

central-western northern African margin as resulting from the evolution of the whole southern Tethyan

margin in response to the Africa-Eurasia divergence.

The Late cretaceous and Paleogene extensions

With the Albian transgression began a sedimentary cycle lasting until the end of Cretaceous in the

platform and until late Eocene in the pre-Atlasic domain. Broadly speaking, during this period the

Source: MNHN. Paris

TECTONIC EVOLUTION OF SOUTHERN TUNISIA

225

Fig. 5.— Late Cretaceous paleostress

evolution in Southern Tunisia.

A, during Vraconian; 1. Vraconian

oucrops; 2. N140 extension; B, during

Cenomanian-Turonian; 1. Cenomanian-

Turonian outcrops; 2, N070 extension;

C, during Senonian; 1, Senonian out¬

crops; 2, N080 extension; 3, multi¬

directional extension.

Fig. 5.— Evolution des paleocontraintes au

Cretace superieur en Tunisie meri¬

dional.

A, au Vraconien ; 1, affleurements

vraconiens ; 2, extension NI40 : B. au

Cenomanien-Turonien ; 1. affleure¬

ments cenomano-turoniens ; 2, exten¬

sion N070 ; C, au Senonien ; I, afleu-

rements senoniens ; 2, extension N080 ;

3, extension multidirectionelle.

Source: MNHN. Paris

226

SAMIR BOUAZIZ ETAL.

tectonic context remains extensional. but the directions of extension did not remain stable and rotated

several times The main stages of this period are illustrated on figure 5. The first stage is the extension

that developed during the Vraconian time (Fig. 5a). In the platform, the Vracoman carbonates are cut by

025-060 normal faults in relationship with a WNW-ESE trending extension. These normal faults are

associated with joints and sedimentary dykes of same direction. Above the Vraconian-Albian deposits,

the Cenomanian sequence exhibits clear syndepositional features along NNW-SSE trending normal

faults They consists in tilted blocks above decollement layers, intraformational breccias and slumps. In

all the sites, scattered from the Gafsa Range in the North, to the platform in the South, the directions of

extension computed from fault-slip data sets are homogeneous in a ENE-WSW direction (Fig. 5b). The

Cenomanian faulting of southern Tunisia may be compared with the extensional structures of the same

age known eastward in the syrt basin and in the eastern Libyan basin. We think that this important

regional extensional event is related to the rifting that started in the African plate during the Early

Cretaceous that separated the African plate in three large blocks (GUIRAUD & MAURIN, 1991). In

Cenomanian, the Tunisian basin, like the Syrt basin, may be considered as the northern extension of the

trans-Saharian basins. According to this interpretation, they should be considered as the consequence of

the general kinematics reorganization linked with the opening of the south Atlantic Ocean (OLIVET et

al., 1984; RlCOU, 1994). " , ,

Above the Turonian Gattar carbonate bar. the Comacian layers display also remarkable

synsedimentary extensional features, well preserved in the Matmata area (BOUAZIZ, 1986; BARRIER et

a/., 1993). The fault tectonic analysis show that the tectonic regime corresponds to a multidirectional

extension with the main directions trending ENE-WSW and NNW-SSE. The ENE-WSW direction is

related to the subsidence of the NW-SE elongated Jeffara basin which started during the Late Cretaceous

(ELLOUZ, 1984; BOLTENHAGEN, 1981). Intra-Cretaceous extensions tectonics have been described in the

Chott Range (ABDELJAOUED & ZARGOUNI, 1981; LOUHAICHI & TLIG. 1993) and in Central Tunisia

(RABHI. 1987). Such a similar extension is known in Central Tunisia where NW-SE ridges and basins

are considered as large scale tilted blocks along 3 major normal faults (BOLTENHAGEN, 1981; BISMUTH

et al., 1982). In Algeria, compressive deformations have been described in the North as early as the end

of Jurassic (OBERT. 1981). They seem to lead to an Albian paroxysmal phase. These compressions are

restricted to the Tellian domain and no trace is visible in southern Tunisia. These deformations are

though to be related to large transcurrent faults accommodating the eastward shift of Africa during the

opening of Atlantic, whereas eastward in Cyrenaica, the Coniacian compressive event (ROHLICH, 1980)

is probably related to the Syrian arc system tectonic.

In the Campanian-Maastrichtian time the directions of extension rotated again from ENE-WSW to

roughly E-W (Fig. 5c). This extension is documented from conjugate normal fault systems that cut the

Campanian carbonates of the platform and the Abiod Formation in the tolded domain. Isopach maps of

Abiod Formation in Sahel and Gabes Gulf indicate a similar NW-SE to WNW-ESE orientation of the

basins (ELLOUZ, 1984). The deformations decrease southward in the platform where deposits are

shallow water carbonates. Northwestward, in Sahel and in the Gabes Gulf, the main Campanian-

Maastrichtian structural directions are NW-SE (HALLER. 1983; ELLOUZ, 1984; OUAL1, 1984; KADRI,

1988).

In Cretaceous, boundary conditions have significantly changed since the Jurassic. The two former

continental masses to the south of Tethys are now separated into four second order plates: (1) America,

Africa, India and Australia-Antartica, and (2) Africa was disrupted into three blocks in the Early

Cretaceous (GUIRAUD & MAURIN, 1991). Geological observations on the Alpine chains from the Alps to

southeastern Asia show the systematic migration of continental blocks from the southern to the northern

side of Tethys and the associated asymmetry of this part of Tethys bounded by passive margins to the

south and active margin to the north. The westernmost block is the Apulian block separated from Africa

probably during the Cretaceous in relation with the opening of the East Mediterranean basin.

In Tunisia, the Mesozoic-Cenozoic boundary does not mark any major tectonic change, while

compressive deformations began earlier in the northern Tethyan margin in front of Apulia considered as

active since the end of Early Cretaceous. In southern Tunisia, during Paleocene the subsidence migrated

northward to the Gafsa basin whereas the stable platform remains emerged until present. In this basin,

no regional unconformity separates the Paleocene deposits from the underlying Late Cretaceous

sequences. Only local unconformities, due to tiltings of blocks have been observed. They are probably

related to the extension tectonic documented from NE-SW normal faults (Fig. 6) and from populations

of normal shear joints. Offshore, the same directions appear in the Paleogene isopachs drawn from

Source: MNHN. Paris

TECTONIC EVOLUTION OF SOUTHERN TUNISIA

227

Fig. 6.— Paleocene-Eocene extensions in the Gafsa Range. 1. Early Cretaceous; 2, Aptian-AIbian; 3. Campanian-

Maastrichtian; 4. Paleocene-Eocene; 5. Mio-PIiocene; 6. Plio-Quatemary; 7. Faults. Diagrams, same than figure 3.

FlG. 6. — Extensions Paleocene-Eocene dans la chaine de Gafsa. I Cretace inferieur ; 2. Aptien-Albien ; 3, Campanien-

Maastrichtien ; 4, Paleocene- ; 5, Mio-PIiocene : 6, Plio-Quaternaire ; 7, failles. Diagrammes, meme legende que

figure 3.

geophysical data (ELLOUZ, 1984). In Central Tunisia, YaIch (1984) describes similar pre-Eocene

normal faults, and in Northwestern Tunisia Zaier (1984) mentioned that the sedimentation was

controlled by NW-SE trending ridges and basins. In the overlying Paleocene-Eocene phosphate layers of

the Metlaoui formation syndepositional tectonic features are commonly observed. They are

synsedimentary normal faults trending 040-060, in relationship with a NW-SE extension. This tectonic

explains the local unconformities observed inside the Metlaoui Formation. A rotation of the o3

directions from NW-SE to N-S occured during Ypresian time after the deposition of the phosphate

layers. Above these layers, the tectonic pattern is characterized by roughly E-W trending normal faults.

Figure 6 shows the paleostress reconstructions in 2 sites of the late Eocene formations of Djebel Alima

illustrating this sub-meridian extension that probably frame the Eocene basin.

The Africa-Eurasia convergence: the cenozoic compressions

The convergence between Africa and Eurasia began in Late Cretaceous. During the Late Cretaceous

and Cenozoic closing phases of Tethys, compressions prevailed on the Peri-Tethyan platform where

they gave rise to intra-plate deformations such as inversion of Mesozoic basins, or wrench faulting. The

timing of such large-scale deformations is diachronic between the northern and southern margin of the

Tethyan Ocean. They started in Late Cretaceous in the European margin whereas the first recorded

compressive event is only late Eocene in age in the African platform.

228

SAMIR BOUAZIZ ET AL.

THE LATE EOCENE DEFORMATIONS: THE ALGERIAN "ATLASIC PHASE"

The Algerian “ Atlasic phase ”, middle to late Eocene, had a major incidence on the entire Alpine

domain (PETIT. 1976; MATTAUER et al., 1977, BOUILLIN, 1986). In Algeria, it is the major tectonic

event responsible for the 030-040 trending folds. In southern Tunisia, this event is poorly documented

and the main compressional tectonic event is late Miocene, in relation with the Tellian nappe tectonics.

In southern Tunisia, and more particularly in the Gafsa range, the Miocene continental deposits

unconformably overly the whole Late Cretaceous to late Eocene sequence. The angle of this regional

unconformity never exceeds few degrees. In consequence, it is very difficult to identify from fault

tectonic analysis, the pre and post-unconformity fault populations, excepted in the case of

synsedimentary tectonics. It should be possible to distinguish a pre-unconformity compression only if

one particular direction of compression, missing in the post-unconformity deposits, exists beneath the

unconformity. It is not the case in the studied area, and all the reconstructed directions of compression

exist in formations located both below and above the unconformity. So, field analysis of fault-slip data

sets does not allowed to identify an eventual late Eocene compressive event. Nevertheless, this regional

late Eocene unconformity pleads for a late Eocene tectonic event. The associated deformations were

probably low and restricted to large scale flexures probably due to reactivation of former faults.

Compressive structures of this age have been described in Central Tunisia (KHESSIBI, 1978; HALLER,

1983). and in northern Tunisia (BISHOP, 1975; BEDIR, 1985; TOUAT1, 1985; BEN AYED, 1986). We may

point out that if this late Eocene compressional tectonic existed, the trend of the compression was

necessarily either very close or similar to the trend of the later Neogene compressions because it does

not appear only in the pre-unconformity formations. In southern Tunisia, the compressive late Eocene

deformations are likely the weak effects of the Atlas folding that marks a new convergent boundary

between stable Africa and Moroccan and Oran Mesetas during this period (LAFFITE, 1939; CAIRE, 1957;

POLVECHE, 1959; GUIRAUD, 1973; OBERT, 1981).

THE MAJOR COMPRESSIONAL PHASE: THE ATLASIC PHASE

With the deposition of the Miocene continental formations started a new cycle of sedimentation

marked by the erosion of the flexures related to the late Eocene tectonics. The brittle tectonic analysis of

these continental formations, mainly composed of sandstones and clays, shows that succeeded during

Neogene and Quaternary: (a) a period of subsidence originating the thick sequences of the Mio-Pliocene

basins, and (b) several compressive events resulting in various types of deformations.

The oldest brittle structures recorded in the Mio-Pliocene formations are NW-SE trending normal

faults populations. Most of them have been reactivated as right lateral strike-slip faults during the later

tectonic events. The Gafsa fault, active during the late Pleistocene, is probably one of these former

major normal faults (Fig. 7). This episode of faulting is generally neglected in the Neogene tectonic

reconstruction, generally focused on the compressive events. The thickness of the Mio-Pliocene

sequence (more than 600 m) as well as its variations of thickness indicate that during this period develop

subsiding basins. Basins of same age and same orientation have been described in northern Tunisia

(ROUVIER, 1977), associated with NW-SE trending synsedimentary normal faults (PHILIP et al., 1986).

In the Jeffara plain, this distension is still active, as observed in Djerba Island (BOUAZIZ, 1986; 1995),

and in the Libyan Jeffara (KEBF.ASY, 1980). These basins are probably of the same origin than the NW-

SE trending Pelagian basins which developed later in Neogene time. It is clear that in the studied area

this distensive event is older than the major phase of folding.

Following this extensional period, came the deformations resulting in the folding of the platform in

the pre-Atlasic domain. The brittle tectonic analysis of the pre-Atlasic domain allowed us to reconstruct

the stress pattern evolution of the compressive events. We identify five groups of directions of o 1 in the

pre-Atlasic domain: NW-SE, 150 to 180, 020-040, 080 to E-W and 100-110 (Figs 7, 8). The two first

directions are the most frequently measured. They are generally sub-perpendicular to the fold axes and

their directions reflect the geometry of folds (Fig. 7). They may be considered as related to the major

phase(s) of deformation. When criteria of relative chronology exist, the oldest compression is the NW-

SE event which occurred generally before folding when the pre-Atlasic domain was still a platform.

Nevertheless, some sites indicate post-tilting NW-SE compressions as well, showing that it is not only a

Source: MNHN. Paris

TECTONIC EVOLUTION OF SOUTHERN TUNISIA

229

pre-folding event. This compression has been recorded both in the stable platform and in the south

Atlasic domain. The so called “ Atlasic phase ” is very regular in both domains as illustrated by the rose

diagram of figure 8.

Neqr me

GAftA]

cHorr fi

^=t>3Q=. i** i EIH ferfil OBI LiiJ

i‘ i 1 1 | | i ‘ i * ^4

.'M & L^L =-- qp

Fig. 7.— Miocene to Quaternary compressions in the South Atlasic domain and relationship between the fold belt and the

compressional axis.

1, NW-SE compression; 2, N-S compression; 3, E-W compression; 4. NE-SW compression; 5. Triassic; 6. Early

Cretaceous; 7, Aptian-Albian; 8. Senonian; 9. Paleocene-Eocene; 10. Mio-Pliocene; 11. Quaternary: 12. Faults.

FIG. 7— Compressions miocene a quaternaire dans le domaine sud-ailasique et relations enlre les plis ei Ies directions de

compressions.

I, Compression NW-SE; 2. Compression N-S; 3. Compression E-W; 4, Compression NE-SW; 5. Trias ; 6. Cretace

inferieur ; 7, Aptien-Albien : 8. Senonien ; 9. Paleocene-Eocene ; 10. Mio-Pliocene ; 11. Quaternaire ; 12, failles.

Fig. 8.— Compressions in southern Tunisia. Rose diagrams of Ol directions computed from fault population analysis, a. in the

whole studied area including the stable platform and the pre-Atlasic folded domain (144 stress tensors); b. in the stable

platform (60 stress tensors); c. in the Southern pre-Atlasic folded domain (84 stress tensors).

Fig. 8.— Compressions en Tunisie meridionale. Diagrammes de frequence des directions de Ol calculees a partir de l'analyse

des populations de failles, a, dans T ensemble de la region etudiee (plate-forme stable et domaine plisst pre-atlasique

(144 tenseurs); b, dans la plate-forme stable (60 tenseurs); c. dans le domaine plisse pre-atlasique (84 tenseurs).

Source

230

SAMIR BOUAZIZ ETAL.

The second direction of compression trends 150-180 (Figs 7, 8). It is particularly well developed in

the folded domain. On the contrary, it is completely missing in the stable platform. The peak of o 1

directions for this sub-meridian extension is wider than the peak of the NW-SE compression (Fig. 8).

These al directions are almost systematically subperpendicular to the strike of the fold where the site of

measure is located. It is illustrated on figure 7 where both NW-SE and NNW-SSE to N-S compressions

are shown. This N-S event is often considered as a major event in southern Tunisia. Nevertheless, its

lack in the stable platform, whereas the NW-SE compression is well printed, may indicate that a large

part of the 150-180 directions of compression reflect more the trends of the fold axes than a major sub¬

meridian regional event. It should mean that the major event, responsible of folding in the southern pre-

Atlasic domain, should be only the NW-SE compression that generated both NE-SW and E-W trending

folds. This geometry may be easily explained by reactivation of former normal faults. The structural

expression of the compression relates to the geometry of the former structures (dip and trend of the

former normal faults) in relation to stress orientation have been modelized by LETOUZEY (1990). In

southern Tunisia where the compression trends obliquely (NW-SE) to the strike of the basin and of the

former major normal faults (E-W) shortening was accommodated by both NE-SW folds and reactivation

of E-W trending Jurassic normal faults. This hypothesis accounts for (a) the wide peak of ol directions

that may reflect the bending of the ranges, and (b) the lack of submeridian compressions in the platform

(Fig. 8). Despite that the major E-W folding does not result probably from a major sub-meridian

compressional event, such a Quaternary compression undoubtedly exits. It results in the right lateral

displacement along the Gafsa fault system which cuts and displaced the folds. This N-S compression

appears on figure 7 around the city of Gafsa where roughly trending N-S ol directions have been

determined from analysis of fault populations in Quaternary conglomerates (Fig. 7). It originates from

the relatively slow (1cm / y) N-S convergence between Africa and Eurasia.

In Central Tunisia, the major event is known to be Tortonian (BLONDEL, 1991) where the folded

Segui formation, Messinian to Villafranchian in age, overlies unconformably the Aquitanian-Tortonian

continental sequence. In the studied domain, such a geometry has not been observed and no

unconformity exists between the Segui formation and the underlying Neogene continental formation as

well as inside the Segui formation. It results that the age of the major deformation phase is at least

subsequent to the beginning of the deposition of the Segui formation and probably post-Tortonian. So, in

southern Tunisia, the age of the main NW-SE trending event appears to be more recent than in Northern

and Central Tunisia where the major deformation phases are respectively dated of Langhian (ROUVIER,

1977) and Tortonian (BLONDEL, 1991), as far as the stratigraphy of the Late Cenozoic continental

sequence is well known. This Late Tertiary tectonic is more a consequence of the north dipping

subduction associated with back-arc opening and the migrations of the western Mediterranean blocks

than the low convergence rate between Africa and Eurasia (TRICARD et al., 1994). The separation of

Corsica and Sardinia from Eurasia initiated in the Late Oligocene and the front of the block collided

with the African margin in the middle Miocene, generating the Tellian Atlas.

The minor neogene-quaternary compressions

Three other compressional events trending 020-040, ENE-WSW to E-W, and 100-110 affected the

studied area (Figs 7, 8, 9). All of them have been observed in the Pliocene-Quaternary formations. In

both the stable platform and the folded domain, there are never associated with large deformations such

as folds or kilometric strike-slip faults.

The ENE-WSW to E-W compression has been determined in both the platform and the folded

domain. In the pre-Atlasic area, the associated fault populations are post-folding conjugate strike-slip

faults. In the Gabes area, strike-slip faults related to this event cut the karst filling of Matmata loams,

indicating at least a middle Pleistocene age.

The 020-040 compression is expressed by populations of strike-slip or reverse conjugate faults. It has

been observed in the whole investigated area. In the folded domain, (a) this compression has been

observed in the Pliocene to Pleistocene Segui formation, and (b) two axes of the reconstructed

paleostress tensors are included in the bedding plane, suggesting a pre-folding event. As the age of the

folding is not clearly defined in this area within the late Miocene to Pleistocene period, a large

uncertainly exists on the age of this compression.

The last compressive event, trending 100-110, has been only observed in the platform (Fig. 8),

essentially in formations ranging in age from Triassic to Senonian. The trends of al related to this event

Source: MNHN, Paris

TECTONIC EVOLUTION OF SOUTHERN TUNISIA

231

FlG. 9.— Tectonic calendar of Southern Tunisia since Late Permian showing the main tectonic events in term of stress

directions.

FlG. 9.— Calendrier tectonique de la Tunisie meridionale depuis le Pennien superieur presentant les principaux evenements

lectoniques en terme de direction de paleocontraintes.

Source: MNHN. Paris

232

SAMIR BOUAZIZ ETAL.

are very regular over the whole studied area of the platform. The faults populations associated to this

event are generally conjugate systems of strike-slip faults. The age of this event is probably Neogene or

Quaternary because this directions have been also computed from fault populations measured in the

Mio-Pliocene marine clays of the Jerba area.

Despite these minor compressions are difficult to assign to a major tectonic event, their large

distribution in the whole investigated area demonstrates that their are not due to local stress fields but

rather to regional events. In addition, they have been described in the Plio-Pleistocene formations of the

northernmost African platform in the Ragusa plateau of southern Sicily (BARRIER, 1992).

CONCLUSION : EVOLUTION OF THE AFRICAN MARGIN IN SOUTHERN TUNISIA

The tectonic calendar shown on figure 9 summarizes the main tectonic events that occurred in

Southern Tunisia since Late Permian. This reconstruction in term of stress direction allows to precise the

regional geodynamical evolution in Tethyan context. During the Mesozoic-Cenozoic, Tunisia occupied

a remarkable position between the Eastern and Western Tethys. Soon in the Permian time before the

break up of Pangea, Southern Tunisia marked the westernmost extension of the Permian rifting

prefiguring the opening between Gondwana and Laurasia. After the Pangea break up, and during Late

Triassic to Early Cretaceous, the tectonic evolution of this area was mainly driven by the opening of the

East Mediterranean basin associated with the north drift of the Apulian block, whereas in the Maghreb

West of Tunisia, developed a transform margin accommodating the rotation between Europe and Africa

associated with the opening of Central Atlantic. So, the southern Tethyan margin in Tunisia marked

during Jurassic the boundary between two tethyan domains: the Atlantic Tethys to the West,

characterized by transcurrent movements, and the westernmost part of Eastern Tethys to the East,

characterized by diverging boundaries (RlCOU, 1994). All the following tectonic history of this region is

controlled by this particular configuration.

In Cretaceous time. East of Tunisia the tethyan margin was a passive margin whereas the

convergence developed far North along the southern Eurasian margin at the front of the Apulian block.

This explains why during Cretaceous E-W to NE-SW extensions dominated and why the compressions

were scarce. In the same period. West of Tunisia, no major active boundary are known and the Mesozoic

plate pattern West of Apulia was governed by a single plate boundary located North of Iberia. This

tectonic pattern lasted until Eocene when collisions developed along the Southern Tethyan margin

forming the Alpine system in North Africa in response to the jump of the Africa-Eurasia boundary, from

North to South of Iberia. The first compressive movements recorded in southern Tunisia are late Eocene.

They are the echoes of the Algerian “ Atlasic folding ”, well documented in Algeria.

In Neogene, after a Miocene NW-SE extension, still active in the Jeffara plain, the major

compressional event occurred between late Miocene and Quaternary, resulting from the collision of

minor blocks migrating southeastward from Eurasia. Since this period, the slow convergence between

Africa and Eurasia appears to be consumed in Western Mediterranean by intra-plate deformations

generating large earthquakes, whereas in Eastern Mediterranean they are consumed in subduction zones.

The Late Quaternary activity of the Gafsa fault probably results from this process. Such a paleostress

reconstruction based on the analysis of many sites provides accurate constraints in geodynamic

reconstructions. This type of study, extended at the scale of the North African margin, may lead to a

better knowledge of the tectonic evolution of the southern tethyan margin from Morocco to Egypt since

Mesozoic.

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Source: MNHN, Paris

Structural inheritance and kinematics of folding and

thrusting along the front of the Eastern Atlas

Mountains (Algeria and Tunisia)

Dominique FRIZON DE LAMOTTE '",Eric MERCIER w ,

Fatima OUTTANI" 1 , Belkacem ADDOUM m , Hacene GHANDRICHE 12 ',

Jamel OUALI 131 , Samir BOUAZIZ 131 & Jean ANDRIEUX 141

(,) Universite de Cergy-Pontoise, Departement des Sciences de la Terre, URA CNRS 1759

95011 Cergy-Pontoise Cedex, France

(2 ' Sonatrach Exploration, Cote Rouge. Alger, Algerie

(3) Ecole Nationale des Ingenieurs de Sfax, Departement de Geologie, BP W. 3038 Sfax, Tunisie

,4> Labratoire de Geologie Structurale, URA CNRS 1369, Universite Paris-Sud, 91405 Orsay Cedex, France

ABSTRACT

On the basis of recent field work, subsurface data from petroleum exploration and geometrical modelling of fold-thrust

structures, this paper describes the structural style and the kinematics of the front of the Atlas Mountains from Biskra (Algeria)

to Tunis (Tunisia). Two successive deformation fronts, superimposed in some places but distinct in other ones, are

distinguished. Both are thrust fronts exhibiting strong virgations from E-W to N-S trends. This quite complex geometry is

linked to the Mesozoic inheritance and to changes in the tectonic transport directions occurring between the two tectonic

compressive events. Everywhere along the fronts and independently from the age of the deformation, fault propagation fold

represents the main folding mode. These folds are frequently altered by late evolutions occurring during the same or distinct

tectonic events. This leads to important changes of the structural style along strike: the front is underlined successively by a

backthrust, a forethrust (blind or emergent) and finally along the N-S segment by a fold cut out by a strike slip fault.

Additionally, jumps in the emplacement of the decollement levels are accommodated by tear faults and small vertical axis

rotations of cover sheets.

RESUME

Heritage structural et cinematique des plissements et chevauchements au front de l'Atlas oriental (Algerie et

Tunisie).

En s'appuyant sur des travaux de terrain et des donnees de subsurface d’origine petroliere et a l'aide de modelisations

geometriques des structures de chevauchement-plissement, cet article decrit le style tectonique et la cinematique du front de la

chaine atlasique depuis Biskra (Algerie) jusqu’a Tunis (Tunisie). Deux fronts de deformations successifs, superposes en certains

endroits, peuvent etre distingues. Les deux montrent une forte virgation en toumani d’une direction E-W a une direction N-S.

FRIZON DE Lamotte, D., Mercier, E., Outtani, F., Addoum, B.. Ghandriche, H., Ouali, J., Bouaziz, S. & ANDRIEUX,

J., 1998.— Structural inheritance and kinematics of folding and thrusting along the front of the Eastern Atlas Mountains

(Algeria and Tunisia). In: S. Crasquin-Soleau & E. Barrier (eds), Peri-Tethys Memoir 3: stratigraphy and evolution of Peri-

Tethyan platforms. Mem. Mus. natn. Hist, nut., 177 : 237-252. Paris ISBN : 2-85653-512-7.

Source: MNHN. Paris

238

DOMINIQUE FRIZON DE LAMOTTE ET AL.

Cette geometric est liee d’une part a l'heritage mesozoique et d’autre part a des changements de la direction du transport

tectonique intervenant entre les deux 6venements compressifs majeurs. Independamment de I'age de la deformation, le mode de

plissement dominant est le pli de propagation de rampe. Ces plis sont souvent affectls d evolutions secondaires survenanl

pendant le meme £venement tectonique ou au cours d'evenements distincts. II en resulte d'importantes variations laterales du

style structural. Ainsi. le front correspond successivement a un retro-chevauchement, un pro-chevauchement (souvent aveugle)

et Finalement, le long des segments N-S, h un decrochement. Par ailleurs, les sauts dans la position des niveaux do decollemcnt

actifs sont accommodes par des failles de ddchirure et de faibles rotations de panneaux de couverture.

INTRODUCTION

In North Africa, three main structural domains are classically considered from South to North

(Fig. 1) the Sahara platform, the Atlas mountains consisting in weakly inverted basins including rigid

elevated regions called High Plateaus or Mesetas and the Rif-Tell chain consisting in strongly inverted

basins involved into nappe structures.

MEDITERRANEAN SEA

Moroccan Mese'

Aures

High Plateaus

Sahara

Alias

0ISKK-4

Sahara plafform

ATLANTIC OCEAN

Basement

Mesozoic-Cenozol'c cover

Folded Mcsozolc-Ccnozoic cover

RIF TELL MOUNTAINS

Basement

Mcsozoic-Ccnozotc cover

}

SAHARA PLATFORM

Fig. 1.— Schematic map of North Africa showing the main structural domains.

Fig. /.— Carte structural schematique de VAfrique du Nord montrant les principaux domaines.

This tectonic framework is partly inherited from breaks which occurred during the Mesozoic.

However, the South Atlas front (Sahara flexure) which is running from Agadir to Tunis (Fig. 1) is not

everywhere superimposed to the paleogeographic transition between the Sahara platform and the Atlas

basin. This is particularly obvious in the eastern area where the front extends far South of the basin limit

and forms a large scale virgation (Fig. 2). Another particularity of this area is the existence of two

successive and no-homoaxial fold-thrust events recognised decades ago (LAFFITTE, 1939; CASTANY,

1949). However the terminology remains quite confuse. Following the more frequent usage, we will use

the terms "Allas" and " Alpine" events to call the first and second event respectively. In order to avoid a

possible confusion between the geographic and the chronological meaning of the term Atlas, we will use

in this paper the typography "Atlas" when we will speak about the "Atlas" event. The problem of the

ages of the two tectonic events will be discussed below. Figure 2 gives a schematic view of the two

successive deformation fronts. The fronts are superimposed along the N-S axis (Tunisia) and West of

Biskra but are distinct between these two regions.

Traditionally, the Folds ot Atlas Mountains have been indistinctly interpreted as resulting from cover

accommodations above basement strike-slip faults (Fig. 3) (AlSSAOUI, 1984; BOUAZ1Z, 1995; FAVRF.&

STAMPFLI, 1992; ViALLY et al., 1994; ZaRGOUNI, 1985). In this paper, we will defend a different

interpretation in which the significance of the two tectonic events are distinguished.

Source: MNHN. Paris

GEOMETRY AND KINEMATICS OF THE FRONT OF ATLAS MOUNTAINS

239

The "Atlas" event led to the uplift (inversion) of the Atlas basin. It is well expressed in the Atlas

Mountains but unknown in the Tell-Rif (or at least in their external zones). The "Atlas" event must

consequently be considered as resulting from an intracontinental evolution which is now better

understood in the frame of the kinematics of the Tethyan region. RlCOU (1994) suggested that after the

collision between Iberia and Africa (45 Ma), the "Atlas" folding marked the new convergent boundary

limiting a "temporary new plate" which comprised Iberia, parts of Atlantic, Western Tethys and a north

African continental stripe including the Moroccan and Algerian Mesetas and the High Plateaus.

The "Alpine" event is, more classically, connected to the closing of the Tethys ocean and the collision

between Africa and continental fragments called Alkapeca (Alboran-Kabylia-Peloritan-Calabria) by

BOUILLIN (1984). These fragments were detached from each other by back-arc spreading resulting likely

Tell Mountains

Triassic-Jurassic basin

Deep Miocene basins

'Atlas" and "Alpine" Folds

Southern limit of the

"Atlas" deformation

Southern limit of the

" Alpine" deformation

Mio-Plio-Quaternary troughs

FlG. 2.— Structural map of the Eastern Atlas Regions showing the geometry of the "Atlas" and "Alpine" deformation fronts.

Note that the fronts are far south of the limit of the Atlas basin.

FlG. 2 .— Carte structural de la partie orientate du domaine atlasique montrant la geometrie des fronts de deformation

"atlasique" et "alpin". Noter que les fronts sont situes loin an sud de la limite du bassin atlasique.

Source:

240

DOMINIQUE FRiZON DE LAMOTTE ET AL.

Fig. 3.— Diagram illustrating the former interpretation in

which the folds are interpreted as cover accom¬

modations above basement strike-slip faults (modified

from Vially et al. , 1994).

FlG. 3— Schema illustrant l'interpretation classique selon

laquelle les plis traduisent utie accommodation de la

couverture au dessus de decrochements de socle

(modifie d'apres VlALLY et al., 1994)

from subduction roll-back and leading to

divergent escapes of continental blocks (ROYDEN,

1993; FRIZON DE LAMOTTE et al., 1990;

LONERGAN & White, 1995). Alkapeca is

constituted of a pile of nappe-complexes

including mantle fragments which forms the

internal zones of the Tell-Rif and Betics. From

the front of the internal zones the thrusting

propagated southward a considerable distance

through the African margin (the present day

external Tell-Rif) then through the Atlas. Along a

N-S Algerian transept from Great Kabylia to the

Aures Mountains, the thrust propagation is

recorded by successive foreland basins that were

progressively incorporated within the orogenic

wedge (GHANDRICHE. 1991). Relatively to the

" Alpine" event the Atlas Mountains may

consequently be interpreted as the foreland fold

and thrust belt of the Tell-Rif. According to this

interpretation, the present day mountain front, the

so-called South Atlas Front, is demonstrated to be

a major and still active thrust front (FRIZON DE

LAMOTTE et al., 1990; JOSSEN & FILALI-MOUTEI,

1992; CREUZOT et al., 1993; OUTTANI et al,

1995).

This paper is based, in particular, on three

recent PhD thesis by GHANDRICHE (1991),

ADDOUM (1995) and OUTTANI (1996). Its aim is to analyse the geometry and kinematics of the South

Atlas regions in the eastern area where the front shows a strong virgation (Fig. 2). This quite complex

geometry is partly linked to the tectonic and paleogeographic inheritance that we will consider firstly.

STRUCTURAL INHERITANCE

Because of its location close to the Atlantic and Tethys oceans, the Meso-Cenozoic history of North

Africa is complex. We will consider successively four periods of important tectonic activity: the

Triassic-Jurassic period, the Mid-Cretaceous period and, finally, the "Atlas" and "Alpine" compressive

periods.

At the beginning of Mesozoic times, breaks affected the northern part of Africa as a response to

opening of both Atlantic and Tethys oceans. More precisely, during the Triassic and lowermost Liassic

the rifting process has produced a series of "en echelon" NE-SW grabens in North Africa (ANDRIEUX et

al, 1989; LAVILLE & PETIT, 1984; WINTERER & HlNZ, 1984). Between the Middle and Late Jurassic,

the oceanic spreading, already active in the Atlantic, began in the Tethyan domain of the Alps

(Lemoine, 1985; Favre & STAMPFLI, 1992; RlCOU, 1994). North Africa was at that time situated on

the transform zone which linked the two ridges and was the site of active extensional-transcurrent

tectonics. Grabens bounded by NE-SW faults and associated basic intrusions are interrupted by E-W

transform zones. The two most important ones were running between Africa and Iberia and within

Africa itself (Fig. 4). They localise the future Tell-Rif on the one hand and High Atlas, Aures, Sahara

Atlas and Tunisia Atlas on the other hand (Fig. 1). Additionally, The Middle Atlas is inherited from a

NE-SW graben where extension was less pronounced.

In the Eastern Maghrebides, Eo-Cretaceous times are characterized by a new increase of subsidence

rate (ViALLY et al., 1994). It seems that subsidence affected more or less uniformably the region situated

east of Algiers meridian. Pre-Cenomanian normal faulting leading to N-S trending tilted blocks is well

expressed in Tunisia along the "axe Nord-Sud" (OlJALI, 1985) and in some other areas of the Tunisia

Atlas. GHANDRICHE (1991) and MARTINEZ (1991), using both geological mapping and subsurface data,

Source:

GEOMETRY AND KINEMATICS OF THE FRONT OF ATLAS MOUNTAINS

241

described in the Aures Mountains listric normal faults branched downward on the Triassic evaporites

and upward on the Cenomanian shale. VILA (1995) has proposed that this normal faulting generated

migration of evaporites leading to the emplacement of large sub-marine salt glaciers. There is no

evidence showing that the Atlas basement is involved in this extensional tectonics. We propose to

connect the cover decollements of the eastern Maghrebides to the nearby Sirte Basin Rifting and its

northward prolongation in the Pelagian sea (BUROLLET & ELLOUZ, 1986) which was active at the same

time. The whole eo-Cretaceous extensional tectonics is sealed by the marine transgression that started in

the middle Albian and reached its maximum in the upper Cenomanian.

Fig. 4.— The basins geometry of North Africa, East of North America, Iberia and Alboran at the Jurassic-Triassic boundary

(from Favre, 1995; modified for the eastern region).

FlG. 4 .— Geometrie des bassins d'Afrique du Nord a la limite Trias-Jurassique (d'apres Favre, 1995 ; modifie pour les parlies

orientates)

Authors have discribed early (i.e. Cretaceous) compressive "phases" in various parts of North Africa

(see for instance OBERT, 1984). This complex question has been already discussed elsewhere (FRIZON

DE LAMOTTE, 1985) and we consider that these phases are only of local significance. According to

VlALLY et at. (1994), we assume consequently that compression of geodynamic significance started

only at Tertiary times. The first tectonic phase, the "Atlas" phase, has been recognised from a long time

in the Aures Mountains (Laffitte, 1939) and in the Hodna basin (GUIRAUD, 1973) and attributed to

middle-late Eocene. In the Aures Mountains, the Atlas phase is reponsible for NE-SW large scale

trending folds. The unconformity showing Oligocene (?) to Miocene marine sediments resting on folded

Mesozoic to Eocene terrains is well expressed at the top of some anticlines (Fig. 5) but appears only as a

cartographic unconformity above the synclines. The second tectonic phase, the "Alpine" phase, is

responsible for E-W trending folds and thrust faults. Generally the "Alpine" structures are superimposed

to earlier "Atlas" folds leading to a complex tectonic pattern "bayonet folds"; VILA (1980). However in

the zone (Fig. 2) where the Alpine deformation Front extends beyond the Atlas Front, the observed

geometry is quite simple

In Tunisia, two successive compressive events showing the same geometric relationships (N140 then

N-S shortening) are also considered (BEN FERJANI et al ., 1990). However the "Atlas" event, attributed to

the middle Miocene (Serravalian to Tortonian) is younger than in Algeria. The "Alpine" event also

242

DOMINIQUE FRIZON DE LAMOTTE ET AL.

Fig. 5.— The unconformity of Miocene marine

deposits on Upper Cretaceous limestone and

the thrusting of the Dj. Chelia anticline onto

the Miocene (from Ghandriche, 1991). This

view is from the South of the Dj. Chelia, see

location on figure 9.

FIG. 5 .— Discordance du Miocene marin sur les cal-

caires du C ret ace Superieur et chevauche-

ment de I'anticlinal du Dj. Chelia sur ce

Miocene (d'apres GHANDRICHE, 1991). La vue,

localisee sur la figure 9, est prise du sud.

appears as diachronic. For instance, in the Timgad area, "Alpine" folds are sealed by a conglomerate of

Villafranchian age whereas southward, close to the South Atlas Front, the same Villafranchian deposits

are involved in the frontal monocline forming the Sahara Flexure.

This suggests that the "Atlas" Deformation front propagated from West to East whereas the "Alpine"

Deformation front propagated from North to South.

SOUTH ATLAS FRONT GEOMETRIC AND KINEMATIC ANALYSIS

METHODOLOGY

The methods extensively used in this paper have been already presented elsewhere (MERCIER, 1992;

OUTTANI et al ., 1995; MERCIER et al., 1997). This section is only a short outline.

In fold-thrust belts one can consider four types of geologically significant fold-thrust interactions

(Fig. 6). One of them (the detachement fold) is developed above a decollement, the others: fault-bend

fold mode I (Suppe, 1983), fault propagation fold (Suppe, 1985; JAMISON, 1987) and "Chester and

Chester" fold (CHESTER & Chester, 1990) are ramp-related folds. For evident geometric reasons, the

development of a detachement fold needs a thick ductile level below the stiff layers affected by

concentric folding. In the studied part of the Atlas Mountains, we have only recognised ramp related

folds characterized by a kink-like style.

The main difference between a fault bend fold (FBF) and a fault propagation fold (FPF) is that in a

FBF the shortening acting at the back of the fold is almost completely transmitted forward whereas in a

FPF it is accomodated within the fold. The Chester and Chester model combines FBF at depth and FPF

in the near suface strata. Generally the surface geometry of the folds is sufficient to discriminate

between FBF and FPF. However and important problem remains in order to define the geometry of

decollements at depth (Fig. 8). To fix the deep geometry of a FBF, three parameters are needed: the

depths to lower an upper detachements and the slip value transmitted forward. The problem is

consequently underdefined. On the contrary, for a FPF the relationships beetween surface geometry and

geometry at depth are unique. Consequently, the occurence of a single FPF in a section provides data

which fix the geometry at depth for the whole section.

However, field structures are generally more complicated that the simple-step models. In particular

secondary evolutions of FPF are frequent and may obscure the basic geometric pattern (JAMISON, 1987;

Suppe & Medwedeff, 1990; MERCIER, 1992; MERCIER et al ., 1997). Two types of evolution known

theoretically have been recognised in the field (Fig. 8): the transport on the upper flat and the

Source:

GEOMETRY AND KINEMATICS OF THE FRONT OF ATLAS MOUNTAINS

243

Fig. 6.— The main fold-thrust interactions leading to the

presently known folding modes.

Fig. 6 .— Les principaux types d‘interactions plis-chevauche-

menis et les modes de plissement associes.

development of a breakthrough thrust. In the Atlas Mountains, the building of the primary fold and the

secondary evolution may be linked to a progressive deformation during a single event or to a renewal of

an "Atlas" structure during the "Alpine" event.

We will examine the geometry an kinematics of the South Atlas Front, from Biskra to Tunis, in three

particular places of the Eastern Maghrebides exhibiting different thrust-fold interactions.

THE FRONT OF THE AlJRES MOUNTAINS (EAST OF BISKRA)

The Aures Mountains expose a thick (up to 10 000 m) sedimentary sequence ranging from Triassic to

Middle Eocene (LAFFITTE, 1939). This sequence is involved in NE-SW trending folds arranged in "en

echelon" pattern (Fig. 9). In the Middle and the North of the Aures, Oligocene continental and Miocene

marine deposits overlie unconformably the folded sequence. The whole pile is involved in E-W trending

folds well exposed in the Timgad basin (GHANDRICHE, 1991) . The interaction between "Atlas" and

"Alpine" events are expressed by different ways. For instance, the Chelia "Atlas" anticline, situated

South of the Timgad Basin (Fig. 9), is cut out by an E-W "Alpine" thrust-fault leading to the

overthrusting of Barremian quartzite onto Miocene deposits (Fig. 10) (GHANDRICHE, 1991). The

Kasserou anticline, situated N-E of the Timgad basin, is also an "Atlas" anticline as indicated by the

unconformity of the Miocene deposits (Fig. 11). However the steep breakthrough thrust fault which

emerges in the core of the fold affects laterally the Miocene and refers consequently to the "Alpine"

244

DOMINIQUE FRIZON DE LAMOTTE ET AL.

Surface geometry

Identification of the folding mode

Fit of the Fault-propagation

fold model ( a single solution)

Fit of the Fault-bend

fold model ( numerous solutions but

only one is compatible with the

50jJ

previous fit)

Fig. 7.— Determination of the geometry at depth of a couple Fault Bend Fold- Fault Propagation fold using the geometric

properties of Fault Propagation Folds (modified from Outtani et al. 1995).

FtG. 7.—Principes de la determination de la geometrie profonde d'un couple pli de cintrage sur rampe- pli de propagation de

rampe en utilisant les proprietes des plis de propagation de rampe (modifie d'apres Outtani et al., 1995)

Transport on

the flat

Deformation

transmitted in

another site

Steep-limb

Breakthrough

Fig. 8.— The two classical alterations of Fault Propagation Folds (modified from Mercier et al., 1997).

Fig. 8.— Les deux evolutions tardives classiques des plis de propagation de rampe (modifie d'apres Mercier et al., 1997).

Source: MNHN. Pans

GEOMETRY AND KINEMATICS OF THE FRONT OF ATLAS MOUNTAINS

245

phase. This thrust-fault keeps an "Allas" orientation but records an important component of sinistral

strike-slip (MERCIER et at ., 1995). Modelling (Fig. 12) is consistent with an interpretation of the

Kasserou as an "Allas" fault-propagation fold altered by a secondary breakthrough fault of "Alpine" age.

Both Chelia and Kasserou anticlines are developed above a Triassic decollement plane and we consider

that the basement is not involved in the Aures folds.

HODNA

TIMGAD

Batna

BASIN

Guerguitt '

\\\\\\''' : ■NVNVSNN'N

Gheheb

Biskra

Neogene deposits

Triassic deposits

/ Main Anticinal

ON"""'"' axes

a y? Main Thrust

— faults

Main Normal

faults

South Atlas Front

Location of the

sections

Fig. 9.— A schematic structural map of the Aur&s Mountains locating the Figures 5. 10, 11 and 13.

FlG. 9 .— Carte schematique des Aures localisant les figures 5, 10, 11 et 13.

FlG. 10.— A cross section through the Dj. Chelia

showing an “ Alpine” thrust cutting through a

former "Atlas" anticline (from Ghandriche,

1991; location on figure 9). Note the

unconformity of the Miocene deposits situated

below the Dj. Chelia thrust. On the other

hand, the Miocene deposits infilling the

Timgad basin rest more or less conformably

on the underlying stratigraphic pile.

FlG. 10.—Coupe a travers le Dj. Chelia montrant un

chevauchement “alpin” recoupant un

anticlinal "atla-sique" (d'apres GHANDRICHE,

1991 : voir la localisation sur la figure 9).

Remarquer la discordance des terrains

miocenes (cf. figure 5). Par ailleurs, le

Miocene situe dans le bassin de Timgad et

situe au toit de Vanticlinal est plus ou moins

concordant sur son substratum.

Southern limit

of the TIMGAD basin

Dj. Chelia

(Chelia "Atlas" anticline)

Dj. Arhane

2 km

" Alpine" fault

J Miocene

□

Cenomanian

] Neocomian

J Senonian -Eocene

Aptian-Albian [

1 Jurassic

] Turonian

Barremian

] Triassic

Along the southern limit of the Aures massif, late Miocene to Lower Quaternary continental

molasses rest conformably on the Meso-Lower Cenozoic succession (i.e. there is no evidence of Atlas

folds). The whole pile (up to Villafranchian) is involved in E-W trending folds constituting the Guerguitt

and Gueheb range (Fig. 9) (LAFFITTE, 1939). The South Atlas Front is underlined by a south-dipping

monocline (about 45°) which is nowhere affected by south-verging thrust-faults. To the South, on the

Source:

246

DOMINIQUE FRIZON DE LAMOTTE ET AL.

Fig. 11.— An over-simplified geological map of the

Dj. Kasserou anticline (modified from Vila

& Guellal, 1977). The Miocene deposits

unconformably overlain an "Atlas" NE-SW

trending fold but are cut out by a break¬

through fault of ‘ Alpine " age.

FtG. II .— Carte geologique tres schematique de

I'anticlinal du Dj. Kasserou (modifie d'apres

Vila & Guellal, 1977). Les depots miocenes

recouvrent en discordance an pli "atlasique"

d'orientation NE-SW mats sont recoupes par

un chevauchement "alpin ”.

NNW

Miocene deposits

SSE

Fig. 12.— A kinematic model of the Kasserou anticline (modified from Mercier et al., 1995).

FlG. 12 .— Modele cinematique de I’anticlinal du Dj. Kasserou (modifie d'apres MERCIER et al., 1995).

Source: MNHN, Paris

GEOMETRY AND KINEMATICS OF THE FRONT OF ATLAS MOUNTAINS

247

Sahara platform, the Miocene to Quaternary sequence is thick and, as a consequence, in the synclines of

the belt the "regional" is elevated by about 1.5 km above the platform. This geometry predates the

development of the fold-thrust belt. It is illustrated on the figure 13 as the result of a vertical fault which

localises the development of a major backthrust leading to the delamination of the sedimentary cover

above the migrating hinge of the frontal monocline. The backthrust is branched on a shallow

decollement situated at the base of Turanian levels. This shallow decollement is necessarily branched

northward on the deep Triassic decollement along which the whole Aures is translated. Along a N-S

section from the Dj. Taktiout to the Sahara flexure, modelling suggests that the shallow decollement is

folded by the Taktiout fault propagation anticline. The branching on the deeper decollement must

consequently be search more to the North. Along a parallel section situated 55 km eastward (Fig. 13), it

appears that a ramp joins directly the two decollements. Such an important variation along strike may

appear unlikely. However one must keep in mind that the frontal structure is oblique to the "Allas" folds

and that, as a consequence, the branching pattern of the thrust faults is certainly complex. Additionally,

it is worth noting that the interpretation proposed for the Taktiout anticline (Fig. 13) is different from a

previous one (FRIZON DE LAMOTTE el al, 1990). The reason is that we had considered only the fault-

bend mode of folding. However, for the purpose of this paper, the difference concerns only the value of

the slip transmitted toward the Sahara platform.

GEOMETRY AND KINEMATICS OF THE SOUTH ATLAS FRONT

IN THE NEGRINE AND METLAOUI AREAS

This area is typically a zone where the “ Alpine ” Deformation Front is far South of the "Atlas"

Deformation Front. Consequently, the folding process may be considered as globally monophase. We

have numerous new data, presented below, in this zone from recent field work by ADDOUM (1995).

The front is constituted by the steep or overturned forelimb of a fault propagation fold. This forelimb

is frequently cross cut by low angle reverse fault (forelimb faults). On the basis of field and subsurface

data, we have already published the results of modelling of ten cross sections (OUTTANI et al., 1995).

Each sections agrees with a shallow decollement situated within Cretaceous levels. The junction of this

shallow decollement with the Triassic decollement of the Aures occurs beneath the Dj. Abiod and Dj.

Onk situated 30 to 60 km North of the present day deformation front. The results are remarkably

consistent within a given section and from one section to another but disagree clearly with the classic

thick skinned concept. To illustrate the kinematic evolution of the front, we have chosen two sections:

on the first one, we can see a combination of a fault bend fold and a fault propagation fold (Fig. 14b); on

the second the two folds are fault propagation folds suffering a transport on the flat and a breakthrough

respectively (Fig. 14a). Both the earlier folds and the secondary evolutions are of "Alpine" age.

The comparison of the shortening ratios from a section to the adjacent one shows a progressive

increase of the shortening within thrust sheet limited by tear faults sub-paralell to the tectonic transport

direction. This model suggests very small vertical axis rotations (about 1.5°) of cover thrust sheets

(OUTTANI etal. 1995).

In the Metlaoui Mountains (Tunisia: Fig. 14b) a small amount of slide (about 1 km) is transmitted

southward to the Chame des Chotts. This slide cannot explain by itself the structural elevation of the

wide "Chaine des Chotts" anticline built on the site of a Jurassic graben (BEN Ferjani et al. .1990). The

excess area balancing method indicates that the wide anticline results from the inversion of the graben

on a deep decollement level. According to field data, the figure 15 suggests that the translation along the

shallow decollement has been stopped by the previously emplaced inverted deep structure.

Geometry and kinematics of the south atlas front along

THE NORTH-SOUTH AXIS AND THE ZAGHOUAN AREA

Along the N-S axis the "Atlas" and "Alpine" fronts are merging in an unique N-S trending anticline.

Additionally, this structure is clearly localised by eo-Cretaceous tilted blocks (OUALI, 1985; MARTINEZ

et al., 1991). This doesn't mean, however, that it results from the direct reactivation of normal faults. On

the contrary it seems that the faulted terranes suffered rigid-body translations above new created ramps.

248

DOMINIQUE FRIZON DE LAMOTTE ET AL.

Fig. 13.— Two cross sections of the South Atlas Front showing the development of the frontal South-Aures back-thrust (see

explanation in the text), a. Section through the Dj. Taktiout anticline; b. Section 50 km East of the Dj. Taktiout (b2) and

proposed pre- “Alpine" restoration. N.B. the vertical attitude of the pre-" alpine" fault is only indicative. An alternative

hypothesis is given in Addoum (1995).

Fig. 13 .— Deux coupes a trovers le front sud-atlasique montrant le developpement du retro-chevauchement sud-auresien (voir

explications dans le texte). a, coupe a trovers I'anticlinal du Dj. Taktiout. b, coupe 50 km a Test, et restauration pre-

alpine proposee. N.B. L'attitude verticale de la faille "pre-alpine ” n'est pas documentee. Une hypothese alternative est

proposee par ADDOUM (1995).

Source: MNHN. Paris

GEOMETRY AND KINEMATICS OF THE FRONT OF ATLAS MOUNTAINS

249

Dj. Bled Er Rmitta

1 km i Redeyef

Dj. Bligi

| Topographic surface

Miocene to Pliocene

Paleocene to Eocene ^

Maastrichian

up. Turanian & Senonian <

Cenomanian & low. Turanian «{

Albian

Cross-Section b

0.89 km

Mandra Fault zones

Dj. Rhifouf

Fig. 14.— Two cross sections through the Negrine (a) and Metlaoui Mountains (b) illustrating the geometry of the front in this

zone.

Note the slide transmitted southward from the Metlaoui mountains $ee explanations in the text).

Fig. J 4 m — Deux coupes a trovers les Monts de Negrine (a) et de Metlaoui (b) illustrant la geometrie du front dans cette zone.

Noter le glissement transmis a I'avant des Monts de Metlaoui (voir explications dans le texte).

N S

Fig. 15.— A sketch diagram without scale

showing how the "Chaine des Chotts"

combined a deep and a shallow decolle-

ments.

FlG. 15 .— Schema de principe sans echelle

montrant la combinaison dans la chaine

des Chotts de decollements profond et

superficiel.

Source

250

DOMINIQUE FRIZON DE LAMOTTE ET AL.

According to other field observations and analogue modelling (VALET, 1990) we emphasise the

difference between localisation and reactivation. Here, clearly, the normal faults localise the subsequent

folding but are not strictly speaking reactivated (CREUZOT et al ., 1992).

The structural pattern is similar to the Aures one but more complex. The examination of angular

unconformities in Neogene to Quaternary deposits indicates that the first fault propagation fold is of

"Arias" age. However, because of the Eo-Cretaceous inheritance, it was originally an oblique fold born

at about 45° of the shortening direction. This fold has been subsequently cross cut by an " Alpine"

breakthrough faults. Slikensides on the faults are almost horizontal showing that they are pure strike slip

faults agreing with a N-S tectonic transport. As in the Aures, modelling shows that the decollement level

is situated within the evaporitic beds of triassic age.

From the N-S axis, the front bends progressively northward up to the Zaghouan area where it trends

NE-SW normal to the "Atlas" shortening. The front is localised by the boundary of the Atlas basin

inherited from the Triassico-Liassic rifting but as in the previous case there is no actual inversion.

Moreover in some places the early normal faults are rigidely transported and preserved at the outcrop

level (CREUZOT et al., 1993).

CONCLUSION

Relatively to the last tectonic event ("Alpine" event) the Atlas Mountains of eastern North Africa

may be considered as the foreland fold and thrust belt of the Tell Atlas. They result from the "collision"

which occured at the end of the Oligocene times and then propagated southward to the South Atlas Front

where the thrusting is still active.

However the Atlas Mountains are not a classic foreland fold and thrust belt (like the Jura for

instance) because the existence of an earlier folding (the "Atlas" event) In our opinion, this "Atlas" event

leading to the inversion of the Atlas basin results from an intracontinental dextral transpresional

movement. At the end of the "Atlas" event, the region was likely more or less comparable to the present

day North Sea. The subsequent oblitaration by the alpine compression led to the present day pattern.

The tectonic framework is, in fact, more complicated because the development of Miocene to

Quaternary grabens trending NW-SE (Fig. 2), cutting the eastern part of the Atlas and leading to the

progressive "disintegration" of North Africa by successive escape of continental blocks like Sicily which

was a part of Africa at least until Tortonian. An important point is that southward thrusting and NE-SW

extension are coeval showing that in this area of the Maghrebides at least two different geodynamic

processes are active at the same time.

ACKNOWLEDGEMENTS

This paper is a contribution to the Peri-Tethys program (project "South Atlas Front"). The order of

the authors is by institution and alphabetical in each instutition. The authors wish to thank P. TRICART

(Universite Joseph Fourier, Grenoble), R. BRACENE and D. BEKKOUCHE (SONATRACH, Algiers) and

J.M. VILA (Universite Paul Sabatier, Toulouse) for helpful discussions and J.-L. MUGNIER and M.

ZAPPATERA for their reviews.

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INDEX

Aalenian 90; 93; 108; 153; 154; 157; 158; 159; 164;

165; 168; 169

Abadeh 83

Acanthocircus diacranocanihos 354

Adeilo 194; 195; 201; 202; 230; 212

Adigrat 195; 202; 203

Aetosaurus 182

Afar 194; 204; 121

Afar Depression 193; 194; 204; 208; 210

Afar rift 212

Afghan block 72

Afghanistan 74; 166

Africa 83; 132; 202; 216; 220; 221; 223; 224; 226;

227; 228; 230; 232; 238; 239; 240; 241

African Craton 187

African plate 216; 224

Agadir 238

Agarmysh Mount 90

Aghdarband 72

Ahl Mado 203

Akchagylian 93; 159

Albian 90; 132; 147; 157; 159; 164; 165; 166; 167;

169; 217; 223; 224; 226; 242

Alboran 239

Alborz 72; 83; 164; 166; 167; 168; 169

Alethopieris subelegans 23; 26

algae 49; 55

Algeria 177; 178; 179; 185; 217; 223; 226; 228; 237;

241

Algerian Meseta 239

Algerian Platform 187

Algerian Sahara 182; 187

Algiers 240

Alima Djebel 219

Alisporites 182

Alkapeca 239; 240

Allal high 186

Alma basin 93

Alma trough 100

Alpine chain 219

Alps 226; 240

Alut 208

Alveosepta jaccardi 201

Amanbulak 82

Amarti 194; 195

Amba Aradam 195

Amerevian 17

America 226

Ammodiscidae 200; 201

ammonite 41; 50; 200; 201; 203

Amoeboceras kitchini ammonite zone 53; 55; 56; 57;

58; 59; 60; 61; 62; 63; 64

Amoeboceras ravni ammonite zone 56

Anatolide 167

Andromeda 54

Anisian 41; 182

Annularia stellata 23

Antalol94; 195; 201; 202; 203; 212

Antartica 226

antelopes 210

Aptian 131; 133; 140; 143; 147; 157; 159; 164; 166;

167; 168; 169; 223

Apulia 226; 232

Apulian block 226; 232

Apuseni 122; 146

Aquitanian 219; 230

Arabia 159; 169; 194; 203; 221

Arabian peninsula 185

Arabian plate 72

Arabic plate 72

Arabic platform 72

Arabo-African plate 146

Archaelithophyllum lamellosum 16

Archaeocenosphaera 55

Arclipeocenosphaera inaequalis 50; 55

Armati 194

Artemovsk 12; 14; 17; 19; 21; 26

Artinskian 27

Ashkhabad fault 82

Asia 226

Assale 200

Asselian 19; 21; 25; 26; 27; 30

Asterotheca arborescens 23

Asterotheca daubreii 23

Asterotheca densifolia 17; 21

"Asterotheca" lamuriana 16; 17

Astrakhan 35

Astroentactinia paronae 49

Astroentaclinia aff. tantilla 49

Astroentactinia cf. crassata 49

Astroentactinia cf. stellata 49

Ataxopliragmidae 201

Atlas 224; 228; 238; 240; 241; 245; 250

Atlas basin 238; 239; 250

Atlas Front 238; 240; 241; 242; 243; 245; 247; 250

Atlas Mountains 237; 238; 240; 242; 243

Atlas Trough 187

Atlasic belt 216

Atlasic domain 216; 217; 220; 229

Atlasic Flexure 178

Aures 240; 243; 245; 247; 250

Aures Mountains 240; 241; 243

254

INDEX

Aur&s Mountains 240; 241; 243

Australia 226

Australian province 55

Autun Basin 23; 26

Autunia conferta 23; 26

Autunian 21; 23; 25; 26; 30

Azov 154

Baborian zone of Tell 178

Bachisarai 108

Badda 208

Badenian 116; 122

Bajocian 20; 153; 157; 158; 159; 165; 166; 168; 169

Bakhmutskaya 26

Balcik 139

Balcik-Vama 143

Balkan 130; 131; 133; 146; 147

Balkanides 146; 147

Baltic Sea 10

Bandar Abbas 72

Barents region 50

Barents Sea 46; 47

Barents-Pechora region 50

Barents-Pechora-Ukhta region 50

Barents-Timan-Pechora Basin 53

Barremian 157; 159; 164; 166; 167; 168; 169

Bashisarai 100

Bashkiria 35; 51

Bashkirian 29; 51

Baskunchak lake 35

Bathonian 93; 154; 157; 159; 164; 165; 168; 203

Bathysiphon sp. 198

Beatrice field 48

Belarus 10

Belaya River 35

belemnite 55; 203

Belogorsk 157

Belogradcik anticline 133

Bentosuchus cf. siishkini 40

Bercovica anticline 133

Berriasian 157; 159; 164; 165; 166; 168

Beshuy 165

Betics 240

Bihen 203

Bihendula 203

Bihor 146

Biscutum dubium 200

Biscutum sp. 198

Biskra 243

Bitak 165

bivalve 39

Black Sea 93; 133; 147; 164; 166; 167; 168; 173

Black Sea Basin 90; 93; 94; 106; 109; 132; 146; 157;

159; 163; 164; 167; 168; 169

Blue Nile gorge 203

Bolshoye Bogdo Mt. 41

Boreal province 50; 54; 55; 68

Boreal region 40

Boreal Sea 42; 43

bovid 210

brachiopod 37; 49; 198; 200

Brachyopoid 182

Brodispora striata 182

Bucliia concentrica zone 53

Buchia 50

Buia 194; 208; 210

Bukobajskaya dipnoan 41

Bulgaria 130; 131; 133; 146; 147; 166

Burgas 133

Buzau 116; 117; 122

Calabria 239

Calamospora sp. 23

Callipterid 23

Callovian 157; 158; 159; 164; 165; 166; 168; 198;

200; 203; 212

Cambrian 50

Camerosporites secatus 182

Campanian 166; 167; 226

Campano-Maastrichtian 219

Capitosaurid 182

Carboniferous 21; 23; 26; 29; 30; 72; 221

Camian 42; 153; 154; 158; 159; 164; 165; 166; 168;

182; 184; 185; 187; 217; 221; 223

Carpathian area 113; 114

Carpathian foothills 113

Carpathian fore-deep 130

Carpathian Mountains 130

Carpathian nappe 115

Carpathian thrust front 115; 116

Carpathians 114; 115; 116; 117; 118; 122; 123; 131;

145; 147

Carpatho-Balkan 129; 146

Carpatho-Balkan chain 130

Casim-Bisoca 116

Caspian Sea 10; 72; 83

Caucasus 94; 99; 100; 154; 159; 163; 164; 168; 169;

173

Caucasus belt 154

Cenomanian 147; 157; 165; 167; 169; 217; 226; 242;

243

Cenosphaera inaequalis 55

Cenozoic 90; 94; 118; 161; 163; 164; 165; 166; 167;

216; 219; 220; 221; 226; 227; 230; 232; 240

Ceratodus bukobaensis 41

Ceratodus gracilis 41

Ceratodus multicristatus 41

Ceratoikiscum 65

Source:

INDEX

255

Ceratoikiscum mertsi sp. nov. 49; 65

Ceratoikiscum spinosiarcuthum 49

Chainc dcs Chotts 247

charophytcs 39; 41

Chatyr Dag plateau 108

Chelia 243; 245

Chemsi Djebel 219

Chott 217; 219; 223

Chroniosaurus dongusensis 39

Chroniosuchus uralensis 39

Circulina 182

Cis-Caspian 42; 43; 48

Cis-Caspian Depression 35; 36; 39; 41; 42; 43

Cis-Ural Trough 35; 36; 37; 39; 40; 41; 42; 43

Cis-Urals 35; 42; 43

Classopolis 180

Columbites zone 41

Columinisporites sp. 21

conchostracan 39; 41; 42

Coniacian 133; 167; 217

conodont 11; 13; 49

Conosphaera sphaeroconus 55

coral 49; 202; 207; 208

Cordaites sp. 16

Corsica 230

Craspedites subditus ammonite zone 56; 58; 60; 61; 62;

63; 64; 65; 66

Crassispora kosankei 23

Crepidolithus sp. 198; 200

Cretaceous 90; 92; 93; 100; 106; 108; 109; 114; 131;

132; 133; 139; 143; 144; 145; 146; 147; 157; 158;

159; 164; 166; 167; 168; 169; 217; 219; 220; 221;

223; 224; 226; 227; 228; 232; 240; 241; 247; 250

Crimea 89; 90; 92; 93; 94; 99; 100; 106; 108; 109;

110; 111; 147; 153; 157; 159; 164; 165; 167; 168

Crimean belt 93; 100; 164

Crimean chain 90

crocodile 210

Crucella 56

Crucella aff. mexicana 50; 56

Crucella crassa 50; 53

Crucella mexicana 56

Crucella squama 50; 56

Cuhnitzschia (al. Lebachia) angustifolia 26

Culmitzschia (al. Lebachia) parvifolia 16

Cyclagelosphaera lacuna 198; 200

Cyclagelosphaera margerelii 198; 200; 201

Cyclagelosphaera wiedmannii 198

Cyclammina sp. 198

Cyphastrea corrugata 207

Cyrenaica 226

Cyrtocalpis hexagonata 62

Dacides 114

Dadoxylon sp. 16

Dagestan 157; 158

Dahar 217; 221; 223

Daixina 17

Daixina enormis 21

Daixina ruzhencevi 21

Daixina sokensis 21

Daixina sokensis biozone 21

Daixina spp. 21

Danakil 193; 194; 195; 203; 208; 210; 212

Danube 144

Darwinula fainae 39

Deinotherium 204; 207; 212

Desset 204; 207; 208

Devonian 10; 46; 47; 48; 49; 50; 52; 58; 59; 68; 72

Dichophyllum cuneata 23

Dictyomitra multipora 63

Dinocephalian 36; 43

dipnoan 41

Dire Dawa 203

disaccates 182

Disacites 182

Discorhabdus (Palaeoponiosphaera) dorsetensis 198;

200

Discorhabdus sp. 201

Djebel Abiod 247

Djebel Alima 226

Djebel Onk 247

Djebel Rehach 223

Djebel Taktiout 247

Djebel Tebaga 223

Djeffara 188

Djerba Island 228

Dnieper Basin 11; 12; 23; 26; 29; 159

Dnieper-Donets 10; 14; 30

Dobrogea 159; 164; 166; 167; 168

Dogali 194; 204; 207; 208; 210

Dogali-Desset 212

Dogali-Massawa 195

Domanik basin 49; 52

Domanik Creek 48; 58; 59; 68

Donets 11; 21

Donets Basin 9; 10; 12; 14; 16; 17; 21; 23; 25; 26; 27;

29; 46; 161; 164

Donguz ichthyofauna 41

Dorbane Basin 185

Dorbane-Dahar Trough 185

Dorogomilovian 17

Dorsoplanites panderi ammonite zone 50; 51; 53; 55

Dragoman 133

Duplicisporites granulatus 182

East Mediterranean basin 216; 221; 226

Eastern Saharan Erg 217

INDEX

256

East-European Craton 85

echinoderm 200; 201

echinoid 198

Edd 194

Egypt 232

El Asker Djebel 219

El Biod 181

El Borma 181

El Agreb - Hassi-Messaoud Horst 185

elephant 210

Eligmus 203

Ellipsovelatisporites 182

Emine 133; 147

Endosporites globiformis 23

Entactinia cf. additiva 49

Entactinosphaera 58

Entaclinosphaera ass idem 49; 59

Entactinosphaera echinata 49; 58

Entactinosphaera grandis 49; 59

Entactinosphaera nigra 49

Entactinosphaera sp. cf. nigra 59

Entactinosphaera variacanthina 49

Enzonalasporites vigens 182

Eocene 90; 94; 99; 106; 108; 109; 133; 144; 145; 146;

147; 153; 157; 158; 159; 164; 165; 166; 167; 169;

173; 219; 227; 228; 230; 243

Epimayaites falcoides 200

Epimayaites sp. 200

Epistomina unzhensis foraminiferal zone 53

Eponides sp. 200

equids 210

Eritrea 193; 194

Eritrean escarpment 194; 208

Eritrean plateau 207; 208; 210

Ernestiodendron filiciformis 26

Eryosuchus 41

Ethiopia 194; 203

Etropole 133

Eucyrtidium haeckeli 62

Eurasia72; 83; 90; 146; 168; 216; 132; 221; 223; 224;

227; 230

Eurasian craton 90

Eurasian plate 72; 83; 146; 147

Europe 130; 146; 159; 169; 223; 232

European margin 226

European platform 90; 146; 153; 173

Evolutinella emeljanzevi - trachammina septentrionalis

zone 54

Excingula 59

Excingula ? hifaria 50; 59

Excingula bifaria 50

Excingula sp. 59

Famennian 52

fish 37; 39

flora 21; 37; 41; 68

Focsani 114; 116

foraminifer 12; 17; 25; 31; 50; 51; 55; 200; 201

Fore-Balkan 131; 133; 137; 144; 146

Foros 106

Foum Tataouine 223

Frasnian 45; 58; 59; 68

Fusulina cylindrica 16

Fusulinella bocki 16

fusulinid 14; 17; 25

Gabeichelti 204

Gabes 217; 226; 230

Gafsa 217; 219; 226; 228; 230; 232

Gardenasporites sp. 23

Gassi-Touil 182

gastropod 198; 201; 208

Gattar 226

Gelendjik 100; 106

Gemanella zone 41

Georgia 167

Getic 115; 122

Ghadames 181; 187; 188

Glomospira sp. 198

Gnathorhiza triassica 40

Gondwana 29; 72; 221; 223; 232

Gorodische 51; 54; 56; 58; 60; 61; 62; 63; 64; 65

Granizospora vulgaris 36

Great Caucasus 90; 147; 153; 154; 157; 158; 159; 164;

165; 167; 168

Great Caucasus-Southern Crimea belt 159; 164; 168

Great Caucasus-Southern Crimea orogen 153; 159

Great Kabylia 239

Gueheb range 245

Guerguitt range 245

Guria Basin 167

Gzhelian 12; 17; 18; 21; 23; 25; 26; 30

Hagiastrum squama 56

Hamada de Tinrhert 185

Haplentactinia 59

Haplentactinia arrhinia 49

Haplentactinia inaudita 49; 53

Haplentactinia rhinophyusa 49

Hassi-Messaoud 180; 181; 186

Hassi-R'Mel 180; 182; 184; 185; 186

Hauterivian 90; 157; 166; 167; 169

Heliosoma echinatum 58

Hercynian 185

Hettangian 154; 159; 164; 165; 166; 168; 182

High Atlas 240

Source:

INDEX

257

High Plateaus 238; 239

hippo 210

Hodna basin 241

Homo 210; 212

Hungary 146

hyaenids 210

Hydrozoa 198; 202; 203

Iberia 232; 239; 240

ichthyofauna 41

Inder lake 35

India 226

Indian 40

Indol 90

Indol-Kuban 94; 106; 153; 159; 164

Ionian Sea 223

Iran 72; 74; 83; 159

Iran plate 83

Ivano-darievka 21

Jablanica line 132

Jamal 83

Jeffara 217; 221; 228; 232

Jigulites 17; 21

Jigulites altus subzone 17

Jigulites jigulensis 21

Jigulites jigulensis biozone 21

Jurassic 45; 46; 48; 49; 50; 51; 52; 55; 56; 57; 58; 59;

60; 61; 62; 63; 64; 65; 71; 72; 74; 83; 90; 92; 93;

100; 131; 132; 153; 157; 158; 159; 164; 165; 166;

168; 169; 173; 193; 200; 217; 219; 220; 223; 224;

230; 240; 247

Kabyle Ridge 178

Kabylia 216; 239; 240

Kalinovo 12; 14; 17; 19

Kalmyksky 154

Kaluga 52

Kara Sea 46

Karabi Yayla plateau 108

Karaganian 159

Karkinitsky 157; 158; 167

Karpinsky Swell 154; 158; 161; 164

Kartanash 16

Kashpir 51

Kasimovian 12; 14; 16; 17; 18; 26; 27; 30

Kasserou 243; 245

Kayasula 154; 159

Kazakhstan 72

Kazanian 36; 38; 39; 42; 43

Kebili Tebaga Djebel 219

Kenya 203

Kerch 94; 99; 100

Kerch Taman 94; 100; 106; 11

Khamovnichian 17

Kiliya-Karkinitsky-Azov belt 159

Kimmerian 72; 74; 100; 111

Kimmeridgian 47; 50; 51; 52; 53; 54; 56; 57; 58; 59;

60; 61; 62; 63; 64; 157; 166; 201; 203; 212

Kizil Kaya 74; 75; 82; 83

Kogva 51

Kolva 51

Komi 48; 55; 56; 57; 58; 59; 60; 61; 62; 63; 64; 68

Kopet-Dagh 72; 74; 83

Kopct-Dagh-Balkan fault 74

Kraiste 146

Krasnovodsk 72

Krevyakinian 16; 17

Kuban 100; 157

Kula 131

Kungurian 27; 42

Kurnubia palasliniensis 198; 200

labyrinthodonts 40

Ladinian 41; 187; 182

Laevigatosporites sp. 23

Langhian 93; 210

Laur^sia 200; 221; 223; 232

Lenticulina biexcavata zone 54

Lenticulina sp. 198; 200; 201

Liassic 191; 217; 220; 223; 242; 250

Libya 185; 217; 221; 223; 224

Libyan basin 226

Liorhynchus 52

Lithocampe 53

Lithocampe cf. terniseriata 52; 63

Lithocampe haeckeli 62

Lithocampe terniseriata 63

Lituolidae 200

Lodevia nicklcsii 23

Lotharingius cf. crucicentralis 198; 200

Lotharingius hauffii 200; 201

Lotharingius velatus 200; 201

Lower Volga 35

Luda Kamcija 131; 133

Lueckisporites virkkia 39

Lugh-Mandera basin 203

Lundbladispora sp. 23

lungfish 36

Maastrichtian 133; 147; 167; 219; 226

Maghreb 223; 224; 232

Maghrebides 178; 240; 241; 243; 250

258

INDEX

mammalian 204

Mangyshlak 72

Manych 154; 158; 159

Marsupipollenites triradiatus 23

Mashad 72

Massawa 194; 204; 207

mastodonsaur tetrapod 41

Mastodonsaurus 41

Matmata 226; 230

Maykop 154; 157; 158; 159; 164; 167

Maykopian 93; 94; 100

Mecsek 146

Medenine 223

Medenine Tebaga 221

Mekele 203

Mekraneb 185; 221

Melanopsis 207; 210

Melekhovian 21

Mendefera Adi Ugri 207

Meotian 116; 117; 118; 120; 123; 199

Mesetas 238; 239

Mesozoic 74; 90; 114; 132; 161; 163; 164; 165; 166;

178; 187; 194; 204; 207; 208; 216; 219; 221; 223;

226; 227; 232; 242

Messinian 230

Metlaoui 247

Middle Atlas 240

Mid-Russian depression 46

Mid Volga bassin 54

Miocene 93; 94; 100; 111; 114; 122; 123; 130; 133;

143; 145; 153; 157; 158; 159; 161; 164; 165; 167;

169; 194; 207; 210; 212; 220; 228; 230; 232; 241;

243; 245; 247; 250

Mio-Pliocene 219; 228; 232

Mir if us us 54

Mitrolithus sp. 198

Moesia 146; 147

Moesian Platform 129; 130; 131; 132; 133; 139; 146;

147; 148; 164; 166; 167; 168

Moesian-Balkan domain 146

Moldavian 117; 118; 120

Moldavides 114; 122

mollusc 207

Monliparus 17

Momiparus monliparus biozone 17

Moroccan Meseta 228

Morocco 223; 224; 232; 239

Moscovian 12; 14; 16; 29

Moscow 12; 13; 17; 18; 21; 22; 26; 29; 43; 52

Mozdok 154; 159

Mt. Assale 204

Mt. Eggerale 204

Myachkovian 16

Namurian 90

nannofossil 55; 68; 198; 200; 201

Nassellarian 50; 54

Nauliloculina oolilhica 198

Negrina 247

Neocomian 217

Neogene 90; 93; 100; 113; 120; 122; 133; 167; 169;

173; 193; 204; 216; 217; 228; 230; 232; 250

Neo-Tethys 72; 188

Neuropteris ovata 17

Nis-Trojan 131; 132

Nodosaria cf. ronda 17

Nogaisk 154

Noginskian 21

Norian 42; 153; 154; 159; 164; 165; 166; 168; 182;

184; 185

North Caucasus 164

North France Basin 10; 17

North Sea 35; 48; 53; 250

Northern France 54

Norway Sea 47

Novo-Fedorovsk rift 154

Nubia 194

NW-Caucasus 89; 90; 93; 94; 99; 100; 106; 108;

109; 11 1

Obsoletes 14

Odessa 93

Odontopteris osmundaeformis 21

Olenekian 40; 41; 42

Oligocene 90; 93; 94; 100; 111; 153; 157; 158; 159;

164; 165; 166; 167; 169; 173; 230; 241; 243; 250

Oran Meseta 228

Orbicella apenninica 207

Orbicella defrancei 207

Orbicella ellisiana 207

Orbicella reussiana 207

Orbiculiforma 57

Orbiculiforma ? retusa 50; 58

Orbiculiforma cf. iniqua 50; 57

Orbiculiforma ex gr. mclaughlini 51; 58

Orbiculiforma iniqua 57

Orbiculiforma mclaughlini 58

Orbiculiforma sp. 58

ostracod 36; 39; 41; 42; 200; 201

Ostrea sp. 207

Oued Mya Basin 181; 185

Ovalipollis pseudoalatus 182

Oxfordian 157; 166; 198; 200; 201; 203

Source:

INDEX

259

Pachelma 52

Pakistan 72

Palaeomutela vjatkensis 39

Palaeomutela umbonata 39

Paleocene 90; 93; 100; 106; 108; 109; 130; 133; 137;

139; 140; 143; 144; 145; 147; 157; 158; 159; 164;

165; 167; 169; 219; 226; 227

Paleodarwinula abunda 36

Paleodarwinula fragiliformis 39

Paleogene 118; 122; 167; 219; 200; 226

Paleoscenidium cladophorum 49

Paleo-Tethys 74

Paleothalomnus sp. 49

Paleozoic 50; 72; 74; 90; 153; 159; 164; 179; 180;

200

Palmatolepis gigas 49

Palmaiolepis punctata zone 42; 43; 49; 50; 58; 59

Palmatolepis subrecta 49

Palmatolepis timanicus 49

Palmyra rift 221

palynoflora 36

palynomorphs 21

Panafrican 180; 185

Pangea 29; 30; 72; 188; 216; 220; 221; 223; 232

Pannonian 146

Pantanellidae 54

Pantanellium 56

Pantanellium sp. A 57

Pantanellium tierrablankaense 50; 56

Paratethys 93; 115; 157

Paraurgonina sp. 201

Pareiasaur-Theriodont fauna 36; 43

Parhabdolithus embergeri nannoplankton zone 54

Paropamisus mountains 74

parotosuchian tetrapod fauna 41

Parotosuchus 41

Parvicingula 50; 51; 54; 55; 60; 63

Parvicingula ? blackhornensis 50; 60

Parvicingula ? cristata 61

Parvicingula ? enormis 50; 61

Parvicingula aff. alata 51; 60

Parvicingula aff. burnsensis 61

Parvicingula aff. haeckeli 52; 62

Parvicingula aff. spinosa 52

Parvicingula aff. thomesensis 63

Parvicingula alata 60

Parvicingula antoshkinae sp. nov. 60

Parvicingula blackhornensis 60

Parvicingula blowi 51; 61

Parvicingula boesii 52

Parvicingula burnsensis 50; 61

Par\’icingula cf. blowi 50; 60

Parvicingula cf. bursnensis 61

Parvicingula conica 50; 51; 61

Par\>icingula cristata 50; 51; 61

Parvicingula elegans 52; 60

Parvicingula ex gr. burnsensis 61

Parvicingula genrietta sp. nov. 62

Parvicingula haeckeli 50; 52; 62

Pan’icingula haeckeli zone 54

Parvicingula hexagonata 51; 52; 62

Parvicingula inornata 50; 52; 63

Parvicingula jonesi 51; 53

Parvicingula khabakovi 51

Parvicingula multipora 51; 52; 63

Parvicingula papulata 50; 63

Parvicingula papulata zone 53

Parvicingula pizhmica 50; 62

Parvicingula pusilla 50

Parvicingula rugosa 50

Parvicingula santabarbarensis 50

Parvicingula simplicima 50

Parvicingula sp. 63

Parvicingula sp. A 60

Parvicingula sp. K 64

Parvicingula spinosa 63

Parvicingula susollaensis 51

Parvicingula thomesensis 63

Parvicingula vera 52

Parvicingula vera zone 53

Parvicingula zyrjanica 51

Parvicingulids 50; 51; 52

Patmasporites densus 182

Pavlovo-Posadian 21

Pay-Khoy 47

Peceneaga-Camena Fault 166

Pechora 45; 19; 50; 53; 54

Pechora River 55; 56; 57; 58; 59; 60; 61; 62; 63; 64

Pechora-Kolva Aulacogen 47

Pecopteris arcuata 23

Pec ten sp. 207

Pelagian basins 228

Pelagian sea 241

pelecypod 198; 200; 201; 208

Peloritan 239

Peltoceras 203

Peltoceras (Unipeltoceras) sp. 198

Peri-Caspian 46

Peri-Caspian Depression 42

Perisphinctes sp. 200

Peri-Tethyan 130; 178; 179; 227

Peri-Tethys 45; 48

Permian 9; 10; 12; 21; 23; 25; 26; 27; 30; 35; 40; 43;

71; 72; 74; 82; 83; 146; 216; 217; 220; 221; 232

Pcrmo-Triassic 35; 74; 220; 221; 223

" Perisphinctes" aff. subcolubrinus 200

"Perisphinctes" orientalis 200

Phormocampe 64

Phonnocampe favosa 51; 64

260

INDEX

Phytosaurus 182

Pityosporites 182

Pizhma Creek 60; 62

P lathy cryphalus ? pumilus 52; 64

Platycryphalus 64

Platycry phalus pumilus 64

Pleistocene 115; 116; 117; 118; 120; 122; 123; 167;

194; 210; 212; 219; 220; 230

Plerastrea cf. profunda 207

Plerastrea sa he liana 207

Pliensbachian 154; 165

Pliocene 100; 115; 116; 117; 130; 132; 133; 1239;

145; 146; 147; 148; 158; 208; 210; 219; 230

Plio-Pleistocene 194; 207; 232

Plio-Quaternary 94; 219

Ploiesti 118

Podorhabdus grassei 198

Polar Urals 47

Polyentactinia circumretia 49

Polyentactinia kosistekensis 49

Poly gnat us timanicum zone 49

Polypodorhabdus spp. 198; 200

Pontian 159

Pontides 93; 147; 159; 164; 166; 167; 168; 169; 173

Porites sp. 207

Potonieisporites novicus 26

Potonieisporites sp. 21; 23

Potonieisporites spp. 23

Praecirculina granifer 182

Praeconocaryomma sphaeroconus 55

Praeconosphaera 56

Praeconosphaera ex gr. sphaeroconus 50; 55

Praeconosphaera sp. 56

Praeconosphaera sphaeroconus 55

Praeobsoletes 16

Praeobsoletes burkemensis biozone 14

pre-Atlasic 217; 224; 228

Pre-Caucasus 153; 154; 157; 168

Pre-Donets 18; 22

Present 216

Preslav 137

Pripyat 10; 46; 52

Professor Ichirkovo 146

Protohaploxypinus spp. 23

Protrit idles 14; 16

Protriticites ovatus - Quasifusulinoides

quasifusulinoides - Praeobsoletes tethydis

biozone 14

Protriticites pseudomontiparus-Obsoletes obsoletus

biozone 14; 16

Pseudocrucella 56

Pseudocrucella aff. prava 56; 41

Pseudocrucella prava 56

Pseudocyclammina sp. 198; 200

Pseudodictyomitrella 63

Pseudodictyomitrella spinosa 63

Punclalisporites confusus 23

Pyrgidan depression 49

Quasifusulina longissima 17

Quasifusulinoides 14

Quaternary 99; 100; 111; 114; 115; 116; 117; 118;

133; 153; 158; 159; 164; 165; 166; 169; 173; 216;

217; 219; 228; 230; 232; 245; 247; 250

radiolaria 49; 50

radiolarian 49; 50; 51; 53; 55; 68; 200; 201

Ragusa plateau 232

Raminer\ f ia mariopteroides 23

Rasht 72

Rechitskian 17

Red Sea 194; 212

Redmondoides. inflatus 198; 200

Redmondoides medius 198; 200

Redmondoides sp. 200

Rehach 221

Reticulosisporites 182

Rhachiphyllum schenkii 23

Rhaetian 42; 154; 159; 164; 165; 166; 184; 185

Rhodope 131; 133; 147; 167

Rhourde el Baguel 182; 185; 186

Rif-Tell chain 238

Riphean 46

Ristola 54

Riyadhella sp. 201

Romania 115; 130; 146

Ruse 145

Russia 10; 12; 48; 55; 56; 57; 58; 59; 60; 61; 62; 63;

64; 68

Russian plate 30; 35; 36; 40; 42; 43; 52; 56; 57; 60;

61; 62; 63; 64; 65

Russian platform 10; 45; 46; 50; 68; 157; 169

Ryazanian 48

Sa Wer 203

Saale Basin 23

Saar Nahe Basin 23

Saccocoma sp. 201

Sahara Atlas 178; 240

Sahara flexure 178; 238; 242; 247

Sahara platform 177; 178; 179; 180; 182; 184; 185;

186; 187; 188; 216; 217; 238; 241; 247

Sahel 204; 212; 223; 226

Sahnites 182

Saint-Etienne Basin 10; 21; 23

" Sakmarella" moelleri biozone 25

Source:

INDEX

261

Sakmarian 25; 26; 27; 30

Salgir Graben 157

Salgir River 90

Salt Range 72

Santonian 133; 167; 217

Saracenaria pravoslavlevi foramini feral zone 53

Saratov 52

Sardinia 230

Sarmatian 93; 100; 111; 115; 116; 18; 122; 123; 139;

140; 158; 159

Schagonella 21

Schagonella minor 21

Schopfipollenites ellipiticus 21

Scythian plate 72; 90

Scythian platform 151; 153; 154; 157; 158; 159; 161;

163; 164; 165; 166; 169; 173

Scythian-Turanian plate 72

Segui 230

Senonian 133; 147; 167; 169; 230

Serbia 133

Serbo-Macedonian 132; 133

serpulid 201

Serravallian 115; 219; 241

Sicily 232; 250

Sidi Stout 217

Siktivkar 60; 62

Silistra 139; 145

Simi Koma 195; 198; 200; 201; 202; 212

Sinemurian 154; 159; 164; 165; 166; 168

Sirte Basin 241

Somalia 194; 198; 203

Sorkh 83

Sorokin 94; 106;

Southern England 39

Spathian 182

Sphaeroschwagerina cf . fusiformis 21

Sphaeroschwagerina constans 25

Sphaeroschwagerina fusiformis biozone 21; 25

Sphaeroschwagerina moelleri biozone 25

Sphaeroschwagerina sphaerica biozone 25

Sphenopleris germanica 23; 26

Spirillinidae 200

sponge 51; 198; 200; 201

spore 11, 182

Sredna Gora 131; 132; 133; 146; 147

Srednegorie 166; 167; 168

Srednogorian rift 133

Stara Planina 131; 133; 144; 146; 147

Staurodictya retusa 58

Stavropol 154; 158

Stephanian 16; 17; 21; 23; 30

Stephanolithion higotii bigotii 200

Stephanolithion bigotii maximum 198

Stichocapsa 54; 65

Stichocapsa ? devorata 51; 65

Stichocapsa aff. devorata 65

Stichocapsa devorata 65

Stichocapsa sp. 65

Stichocapsa sp. A 65

Stichocapsa sp. B 65

Strandza 146; 147

Striatopodocarpites 39

Striatopodocarpites tojmensis 37

Subcarpathian nappe 116

Suchonella typica 39

Suchonellina fragiloides 39

Suchonellina futschiki 39

Sudak 106

suids 210

Svoge anticline 133

Syrian arc 226

Syrt basin 226

Sysola 45; 51

Syssola Basin 51; 54

Taktiout 247

Talemzane 180; 182; 186

Talesh 72

Taman peninsula 94

Tambov-Tula 52

Tarkhanian-Konka 154

Tastqanites 49

Tataouine basin 217; 223

Tatarian 25; 27; 30; 36; 39; 42; 43

Tauride 169

Tell Atlas 250

Tellian Atlas 230

Tellian Atlasic chain 217

Tellian domain 219; 226

Tellian nappes 216; 228

Tell-Rif 239; 240

Terek-Caspian 153; 158; 159; 164

Tertiary 10; 93; 94; 122; 131; 146; 204; 230; 242

Tethyan Ocean 187; 217; 220; 227; 240

Tethyan Realm 55

Tethys 29; 36; 43; 90; 93; 188; 221; 223; 226; 227;

232; 239

tetrapod 36; 39; 40

Tetrentactinia cf. gracilispinosa 49

Thalassinoides 198

Tigrai 203

Timan 52

Timan-Pechora 46; 47; 48; 49; 50; 52; 54

Timgad basin 242; 243

Tirolites cassianus 41

Tisza 122; 146; 147

Tithonian 157; 159; 164

Tizi-n-Test 223

Toarcian 154; 165; 168

262

INDEX

Tortonian 116; 122; 219; 230; 241

Transcaucasus 159; 164; 166; 167; 168; 169

Trans-Urals 42

Trematosaurid 182

Triassic 25; 30; 36; 39; 41; 42; 43; 72; 82; 83; 90; 93

153; 154; 158; 159; 164; 165; 166; 168; 173; 177

178; 179; 180; 181; 182; 184; 185; 187; 188; 217

220; 221; 223; 230; 232; 240; 241; 243; 247; 250

Tripartites aductus 23

Triticiles 17

Triticiles acutus 17

Triticites acutus-Triticites qucisiarcticus biozone 17

Triticiles quasiarclicus 17

Triticites rossicus 15

Triticites rossicus-Triticites stuckenbergi biozone 17

Tschokrakian 159

Tsugaepollenites oriens 180

Tuapse 94; 100; 106; 167

Tuarkyr 72; 74

Tubiphytes 16

Tulcea basin 159

Tunis 238; 243

Tunisia 182; 185; 188; 215; 216; 217; 219; 220; 223;

224; 226; 228; 230; 232; 237; 238; 240

Tunisia Atlas 240

Tupilakosaurus sp. 40

Turan 159

Turan plate 71; 72; 74; 83; 159

Turan platform 157

Turkmenbasi 72; 74; 75; 76; 82; 83

Turkmenia 166

Turkmenistan 72; 83

Turonian 133; 167; 169; 217; 219; 226; 250

Ufimian 36; 42

Ufra 72; 74; 76; 82

Ukhta 48; 49; 50; 51; 56; 57; 58; 59; 60; 61; 62; 63;

64; 68

Ukraine 10; 11

Ukrainian coal-bearing zone 10

Ukrainian crystalline massif 10

Ukrainian horst 30

Ukrainian Shield 161; 164

Ultradaixina bosbytauensis - Daixina robusta

biozone 21

Upper Thrace 133

Ural 11; 25; 26; 27; 35; 42

Urengoy depression 49

Urzhum Platform 39

USA 30

Valanginian 90; 133; 137; 140; 143; 166

Valea Lunga 116; 117; 118; 122

Vallasporites ignacii 182

Vama block 166

Veliko Tamovo 144

Verrucosisporites 182

vertebrate 36; 182

Vidin 140

Villafranchian 219; 230; 241

Villany 146

Vittatina 26

Volga 35; 40; 46; 54

Volga-Ural 35; 36; 42; 45; 46; 48; 52; 54

Volgian 47; 49; 51; 52; 54; 55; 56; 58; 60; 61; 62; 63;

64; 65

Voronezh 10; 29; 30

Vraconian 217; 226

Vyatka 46

Walchia fal. Lebachia) piniformis 16

Wallachian 115; 116; 117; 118; 122

Watznaueria 198; 200

Watznaueria barnesae 198; 200; 201

Watznaueria britannica 198; 200; 201

Watznaueria communis 198; 200; 201

Watznaueria communis nannofossil zone 200

Watznaueria contracta 200

Watznaueria fossacincta 198

Watznaueria manivitae 198; 200; 201

Watznaueria sp. 201

Watznaueriaceae 200

West Kuban 154

western Europe 11; 30

Westphalian 16; 17; 90

Wetlugasaurus angustifrons 40

Xiphosphaera echinatum 53

Yaaila 157

Yalta 106

Yapetus 48; 49; 52; 68

Yarenian ichthyofauna 41

Yausian 117

Ypresian 219; 227

Zaghouan 250

Zarzaitine 182; 185; 187

Zeugrhabdotus erect us 198; 200

Zeugrhabdotus spp. 198; 200

Zula Gulf 207

Rkmerciemf.nts aux rapporteurs / Acknowledgements to referees

La Redaction tient a remercier les experts exterieurs au Museum national d’Histoire naturelle dont les norns suivent, d’avoir

bien voulu contribuer, avec les rapporteurs de rEtablissement, a revaluation des manuscrits (1995/1997) :

The Editorial Board acknowledges with thanks the following referees who, with Museum referees, have reviewed papers

submitted to the Memoires du Museum (1994/1997):

Afzelius B.

Stockholm

Suede

Lemaitre R.

Washington

USA

Akam M.

Cambridge

Grande-Bretagne

Lovelock P. E. R.

La Haye

Pays-Bas

Andersen N.

Copenhague

Danemark

Machida Y.

Kochi

Japon

Baba K.

Kumamoto

Japon

MacKinnon D.

Christchurch

Nouvelle-Zelande

Bachmann G. H.

Halle-Wittenberg

Allemagne

MacPherson E.

Barcelona

Espagne

Bally A. W.

Houston

USA

Maddison D.

Tucson

USA

Bellido A.

Paimpont

France

Manning R.

Washington

USA

Berggren M.

Fiskebackskil

Su&de

Markle D.

Oregon

USA

Bernoulli D.

Zurich

Suisse

Masakj S.

Hirosaki

Japon

Bertotti G.

Amsterdam

Pays-Bas

Mascle A.

Rueil-Malmaison

France

Bessereau G.

Rueil-Malmaison

France

Mauchline J.

Oban

Grande-Bretagne

BestM.

Leiden

Pays-Bas

McLaughlin P.

Washington

USA

Bourseau J.P.

Villeurbanne

France

McLennan D.

Toronto

Canada

Bruce J.

Helensvale

Australie

Merrett N.

Londres

Grande-Bretagne

BruceN.

Copenhague

Danemark

Messing C.

Dania

USA

Brunton H.

Londres

Grande-Bretagne

Morand S.

Perpignan

France

Carpenter J.

New York

USA

Nakamura L

Kyoto

Japon

Cassagneau P.

Toulouse

France

Naumann C.

Bonn

Allemagne

Chace F. A.

Washington

USA

Newman W. A.

San Diego

USA

Child C. A.

Washington

USA

Oliva R.

Barcelone

Espagne

Cherix D

Lausanne

Suisse

OroussetJ.

Paris

France

ClobertJ.

Paris

France

Packer L.

York

Canada

Cloetingh S.

Amsterdam

Pays-Bas

Peter NG

Singapore

Singapour

Cohen D.

Los Angeles

USA

Plateaux C.

Nancy

France

CookP. L.

Victoria

Australie

Poccia D. L.

Amherst

USA

CORNUDELLA L.

Barcelone

Espagne

Poore G.

Victoria

Australie

CUZIN-ROUDY J.

Villefranche / Mer

France

Raikova 0.

Saint-P6tersbourg

Russie

Darlu P.

Paris

France

RentzD. C. R.

Canberra

Australie

Danchin E.

Paris

France

RichardsAV.

Miami

USA

Davie P.

Brisbane

Australie

Roberts C.

Wellington

Nouvelle-Zelande

Dejean A.

Villetaneuse

France

Roure F.

Rueil-Malmaison

France

Deleporte P.

Paimpont

France

Salomon M.

Marseille

France

Dietrich C.

Champaign

USA

Sazonov Y.

Moscou

Russie

Duffels J. P.

Amsterdam

Pays-Bas

ScholtzC.

Pretoria

Afrique du Sud

Eldredge L. L.

Hawaii

USA

SchmidS. M.

Bale

Suisse

Fahay M.

Highlands

USA

SchwanderM.

La Haye

Pays-Bas

Fleury A.

Orsay

France

Spiridonov V.

Moscou

Russie

Fransen C.

Leiden

Pays-Bas

Stefanescu M. 0.

Bucarest

Roumanie

Gagne R.

Washington

USA

Stewart A.

Wellington

Nouvelle-Zelande

Gorin G.

Geneve

Suisse

Takeda M.

Tokyo

Japon

Guglielmo L.

Messina

Italie

TanC. G.S.

Singapore

Singapour

Gullan P.

Canberra

Australie

Tassy P.

Paris

France

Gunzenhausf.r B.

Zurich

Suisse

Thorne B.

Maryland

USA

Harmelin J.G.

Marseille

France

Tudge C.

Brisbane

Australie

Healy J.

Brisbane

Australie

Van Ameron H. W. J.

Krefeld

Allemagne

Heemstra P.

Grahamstown

Afrique du Sud

Van Baaren J.

Rennes

France

Hodgson C.

Ashford

Grande-Bretagne

Vernon P.

Paimpont

France

HolthuisL. B.

Leiden

Pays-Bas

Vickery Vernon R.

Ste-Anne / Bellevue

Canada

Holthuis L. H.

Leiden

Pays-Bas

Vul M. A.

Lvov

Ukraine

Horvath F.

Budapest

Hongrie

WageleJ. W.

Bielefeld

Allemagne

Ingrisch S.

Frankfurt

Allemagne

Watson N.

Armidale

Australie

Jordan P.

Solothum

Suisse

Wenzel J.

Colombus

USA

Kensley B.

Washington

USA

WiegmannB.

Maryland

USA

Kielan-Jaworowska Z.

Oslo

Norvege

Wilson M.

Cardiff

Grande-Bretagne

Komai T.

Chiba

Japon

WilsonS.

Warrensburg

USA

Krapp F.

Bonn

Allemagne

Yeates D.

Brisbane

Australie

Kristensen N.

Copenhague

Danemark

Young P.S.

Rio de Janeiro

Brasil

Lagardere J.P.

La Rochelle

France

Zezina 0.

Moscou

Russie

Laubscher H.P.

Bale

Suisse

Ziegler P. A.

Bale

Suisse

ACIIEVE iVIMPRIMER

EN MAI 1998

SI R LES PRESSES

L IMPRIMERIE F. PAILLART

A ABBEVILLE

Date de distribution : 15 mai 1998.

Depot ttgal: Mai 1998

N° d’impression : 10341

Source: MNHN. Paris

1 9 MAI 1998

Source: MNHN. Paris

DERNIERS TITRES PARUS

RECENTLY PUBLISHED MEMOIRS

A partir de 1993 (Tome 155), les Memoires elu Museum sont publies sans indication de serie.

From 1993 (Volume 155), the Memoires du Museum are published without serial titles.

Tome 176 : Alain CROSNIER (ed.), 1997. — Resultats des Campagnes MUSORSTOM. Volume 18.

570 pp. (ISBN : 2-85653-511-9) 600 FF.

Tome 175 : Isabel PEREZ FARFANTE & Brian KENSLEY, 1997.— Penaeoid and Sergestoid Shrimps

and Prawns of the World: Keys and Diagnoses for the Families and Genera. 233 pp.

(ISBN : 2-85653-510-0) 350 FF.

Tome 174 : Bernard SERET (ed.), 1997.— Resultats des Campagnes MUSORSTOM. Volume 17.

213 pp. (ISBN : 2-85653-500-3) 245,04 FF.

Tome 173 : Philippe GRANCOLAS (ed.), 1997.— The Origin of Biodiversity in Insects:

Phylogenetic Tests of Evolutionary Scenarios. 356 pp. (ISBN: 2-85653-508-9) 490 FF.

Tome 172 : Alain CROSNIER & Philippe BOUCHET (eds), 1997. — Resultats des Campagnes

MUSORSTOM. Volume 16. 667 pp. (ISBN : 2-85653-506-2) 612,60 FF.

Tome 171 : Judith NAJT & Loic MATILE (eds), 1997. — Zoologia Neocaledonica, Volume 4.

400 pp. (ISBN : 2-85653-505-04) 450 FF.

Tome 170 : Peter A. Ziegler & Frank Horvath (eds), 1996. — Peri-Tethys Memoir 2: Structure

and Prospects of Alpine Basins and Forelands. 552 pp. + atlas. (ISBN : 2-85653-507-0)

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Tome 168 : Alain CROSNIER (ed.), 1996. — Resultats des Campagnes MUSORSTOM. Volume 15.

539 pp. (ISBN : 2-85653-501-1) 550 FF.

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654 pp. (ISBN : 2-85653-217-9) 612,50 FF.

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The Peri-Tethys Programme, started in 1993, examines the influence of Tethyan evolution on

the bordering cratons since its birth (through the break-up of Pangea), its life (by the

extension and formation of oceanic seaways) and finally its death (by collision between the

main bordering plates which led to inversion within the epicratonic basins).

The Peri-Tethys Memoir 3 is subdivided into two parts which correspond to the two great

geographic areas involved in the Program: the Northern platform (9 papers) and Southern

platform (4 papers). The first phase of the Program is based on the research in the eastern

countries and particularly in the New Independent States. This is reflected in the contents of this

volume. The first four papers concern stratigraphy and correlation problems in Ukraine and

Russia from the Upper Carboniferous to Jurassic. The next three are on paleostress problems of

the Crimea- Caucasus and Moesia. The final two on the Northern platform present the tectonic

history of the Scythian platform and Black Sea region. For the Southern platform, the first

paper describes the Triassic series on the Saharan Platform in Algeria (Peri-Tethyan onlaps and

related structuration). The topics of the second paper are the Jurassic and Neogene

sedimentation in Eritrea; the last two papers deal with paleostress evolution in Tunisia since the

Permian: and the structural inherence and kinematics of folding and thrusting along the front of

the eastern Atlas Mountains (Algeria and Tunisia).

Sylvie CRASQUIN-SOLEAU and Eric BARRIER (CNRS, Universite Pierre et Marie Curie. Paris)

coordinated this volume which arose from the International Peri-Tethys Meeting held in Milano

(June 1995).

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