Pi, &■& j

r r Peri- Tethys Memoir 2

Structure and Prospects

of Alpine Basins and Forelands

Edited by

Peter A. Ziegler

& Frank HorvAth

TEXTS

MEMOIRES DU MUSEUM NATIONAL D HISTOIRE NATURELLE

TOME 170

1996

Source : MNHN. Paris

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Cover illustration:

The Alpine orogen. Block diagram outlining the role of strain-partitioning in areas of oblique convergence or collision

(see paper by Ziegler & Roure, p. 21). 6

Source : MNHN. Paris

Peri-Tethjs Memoir 2

Structure and Prospects

of Alpine Basins and Forelands

Source : MNHN. Paris

ISBN : 2-85653-507-0

ISSN : .1243-4442

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MEMOIRES DU MUSEUM NATIONAL D’HISTOIRE NATURELLE

TOME 170

GEOLOGIE

Peri-Tethys Memoir 2

Structure and Prospects

of Alpine Basins and Forelands

edited by

Peter A. ZIEGLER * & Frank HORVATH **

Geological- Paleontological Institute

University of Basel

Bernoullistrasse 32

CH-4056 Basel

Switzerland

** Department of Geophysics

Eotvos .University

Ludovika ter 2

H-1083 Budapest

Hungary

v-A. *

6.B -,u |

vv

EDITIONS

DU MUSEUM

PARIS

1996

Source : MNHN, Paris

Source : MNHN , Paris

CONTENTS

Pages

Preface

J. Dercourt & M. Gaetani . 1 1

Foreword . 13

P. A. Ziegler & F. Horvath

Architecture and petroleum systems of the Alpine orogen and associated basins . 15

P. A. Ziegler & F. Roure

Tectonic evolution and paleogeographv of Europe . 47

Enclosures 1-13

P. O. Yilmaz, I. O. Norton, D. Leary & R. J. Chuchla

Accretion and extensional collapse of the external Western Rif (Northern Morocco) . 61

Enclosures 1-2

J. F. Flinch

Triassic- Jurassic extension and Alpine inversion in Northern Morocco . 87

Enclosures 1-4

M. Zizi

The Valencia Trough: geological and geophysical constraints on basin formation models . 103

M. Torne, E. Banda & M. Fernandez

Geodynamics of the Gulf of Lions: implications for petroleum exploration . 129

R. VlALLY & P. TREMOLIERES

The Aquitaine Basin: oil and gas production in the foreland of the Pyrenean fold-and-thrust

belt. New exploration perspectives . 159

Enclosures 1-6

M. Le Vot, J. J. Biteau & J. M. Masset

Cenozoic inversion structures in the foreland of the Pyrenees and Alps . 173

F. Roure & B. Colletta

8

CONTENTS

Structure and evolution of the Central Alps and their northern and southern foreland basins .. 211

Enclosure 1

P. A. Ziegler, S. M. Schmid, A. Pfiffner & G. Schonborn

The Jura fold-and-thrust belt: a kinematic model based on map-balancing . 235

Enclosures 1-2

Y. Philippe, B. Colletta, E. Deville & A. Mascle

Evolution, structure and petroleum geology of the German Molasse Basin . 263

Enclosures 1-4

D. Roeder & G. Bachmann

Hydrocarbon exploration in the Austrian Alps . 285

Enclosure 1

W. Zimmer & G. Wessely

Hydrocarbon habitat of the Paleogene Nesvacilka Trough, Carpathian foreland basin,

Czech Republic . 305

J. Brzobohaty, S. Benada, J. Berka & J. Rehanek

Development and hydrocarbon potential of the Central Carpathian Paleogen Basin, West Car¬

pathians, Slovak Republic . 321

M. Nemcok, J. F. Keith, Jr & D. G. Neese

Structure and hydrocarbon habitat of the Polish Carpathians . 343

Enclosures 1-3

G. Bessereau, F. Roure, A. Kontarba, J. Kusmierek & W. Strzetelski

Oil and gas accumulations in the Late Jurassic reefal complex of the West Ukrainian Car¬

pathian foredeep . 375

T. S. Izotova & I. V. Popadyuk

New data on the structure and hydrocarbon prospects of the LIkrainian Carpathians and

their foreland . 391

Ya. V. Sovchik t & M. A. Vul

Tectonic setting and hydrocarbon habitat of the Romanian external Carpathians . 403

O. Dicea

Do hydrocarbon prospects still exist in the East-Carpathian Cretaceous flysch nappes ? . 427

Enclosure 1

M. Stefanescu & N. Baltes

Neoalpine tectonics of the Danube Basin (NW Pannonian Basin, Hungary) . 439

Enclosures 1-3

G. Tari

Source : MNHN. Paris

PERI-TETHYS MEMOIR 2: ALPINE BASINS AND FORELANDS

9

Structural-stratigraphic evolution of Italy and its petroleum systems . 455

Enclosures 1-3

L. Anelli, L. Mattavelli & M. Pieri

Relationship between tectonic zones of the Albanides, based on results of geophysical studies . 485

A. Frasheri, P. Nishani, S. Bushati & A. Hyseni

Crimean orogen : a nappe interpretation . 513

I. V. Popadyuk & S. E. Smirnov

3D geometry and kinematics of the N. V. Turkse Shell thrustbelt oil fields, Southeast Turkey . 525

N. Gilmour & G. Makel

Source : MNHN, Paris

Source : MNHN , Paris

PREFACE/ PREFACE

La Tethys s’est installee au sein de la Pang6e & partir

du Trias. On la retrouve aujourd'hui emigre ou en

lambeaux cn Amerique centrale, dans l’Atlantique, dans

les chatnes mediterran6ennes ct moyennes orientales,

L Himalaya et lcs montagnes indonesiennes. Elle debute

par des distensions permo-triasiques affectant de vastes

surfaces pangeennes, puis L accretion se localise au cceur

de ce domaine distendu et ouvre une voie d’eau au

Jurassique superieur entre la Laurasia et le Gondwana. Le

plus souvent, la distension entre ces megacontinents se

fait h la faveur d’accretions oc£aniques generatrices de

plaques peu nombreuses aux limites simples, a l’image de

l’Atlantique actuel. Ulterieurement, des collisions, des

subductions, des obductions font progressivement

disparattre cet ocean, hormis dans l’Atlantique entre

Afrique du Nord-Ouest et Etats-Unis ou il perdure.

On examine dans ce volume une partie seulement du

domaine tethysien, celle comprise entre l’Europe d’une

part et V Afrique, l’Arabie et le Moyen-Orient, d’autre

part. La, la complexity est forte car les plaques

europeennes et africaines ne se separent pas

simplement ; des microcratons s’egrenent entre les deux

megacratons ; de nombreuses plaques existent, ce qui

cree un seuil lithospherique mediterraneen qui, lors de la

collision des megacontinents egeens. cree les chatnes

fort complexes analysees dans ce m6moire. De tels seuils

sont connus & l’ouest et a Test de la Tethys dans des

domaines limits, ce sont le seuil des Caraibes entre les

Ameriques et le seuil indonesien entre les bassins

oceaniques, pacifiques et indiens.

Les chatnes pr£seni£es dans ce memoire sont

analysees a partir des domaines gcologiques de surface et

de subsurface obtenus en usant de techniques el de

methodes recentes et des concepts actuels, mais aussi &

partir de donn6es g£ophysiques nombreuses ; la plupart

d’entre elles avaient jusqu'k present valeur patrimoniale

pour les entreprises industrielles qui les avaient acquises,

d’autres ont ete tres recemment acquises par des

programmes internationaux. Peter A. Ziegler el Frank

HorvAth ont coordonne des leur conception ces di verses

monographies sans jamais contraindre les auteurs a se

conformer a un cadre inutilement rigide. Les bassins

llexuraux k l’avant des chatnes ct les bassins molassiques

<i Larriere des chatnes sont particulierement etudies et

renouveles ; chaque lecteur peut lui-meme £tablir de

fructueuses comparaisons.

The Tethys Ocean set itself in place at the centre of

Pangea from the Trias. Remnants can be found even today

either complete or in outliers in Central America, in the

Atlantic, in the Mediterranean and Middle Eastern ranges,

the Himalayas and the mountains of Indonesia. It began

when Permo-Triassic distension affected vast areas of

Pangea, followed by accretion concentrated at the heart of

this extended region and. in the Upper Trias sic, opened

up a limb of water betv.'een Laurasia and Gondwana. In the

most Jrec/uent of cases, distension between these two

supercontinents occurred with the aid of oceanic

accretions that generated a small number of plates with

straightforward boundaries, like the present Atlantic.

Later on collisions, subductions and obductions

progressively closed up this ocean, obliterated except in

parts of the Atlantic between North-West Africa and the

USA where it persists.

This book focuses on just one part of the Tethyan

realm between Europe on the one hand and Africa, Arabia

and the Middle East on the other. The situation here is

highly complicated because the boundary between two

major plates - the European and African - is far from

simple. Strings of small plates formed between the two

super-cratons to constitute a Mediterranean lithospheric

sill which at the time of the Aegean collision of the

supercontinents created the extremely complex mountain

belts analyzed in this work. Such sills are known in

restricted areas to the west and east of Tethys: the

Caribbean sill between North and South America and the

Indonesian sill between the great basins of the Pacific

and Indian Oceans.

The mountain chains featured in this work are

examined from the starting point of surface and

subsurface geological domains found using modern

techniques and methods and current concepts, and also

from a large amount of geophysical data, most of which

up to the present formed part of the heritage of the

industrial companies that had acquired them ; other data

have been acquired very recently through international

research programmes. Peter A. ZlEGLER and Frank Horvath

have coordinated these diverse monographs since their

inception without restricting the authors to compliance

with an unnecessarily strict framework. Particular focus is

placed on updating notions on flexural basins in front of

the chains and molasse basins behind the chains: readers

can hence make their own fruitful comparisons.

12

PREFACE

Les chaines peri-mdditerraneennes et les chaines

associees r^sultent de Involution de talus et de marges

tethysiennes ; la genesc des bassins flexuraux dans les

avant-pays est une reponse des cratons et des

microcratons du seuil lithospherique mSditerrancen a

mpaississement crustal.

Les distensions ant6-t6thysiennes, celles assocides &

F accretion, se font sentir loin sur les cratons bordiers

bien au-del<t des domaines qui s’infl^chiront sous les

chaines et deviendront des bassins flexuraux. Ensuite,

lors des Stapes de la collision des megacratons et des

microcratons constitutifs du seuil lithospherique

mediterraneen, des subsidences reprennent le long de

failles fragiles et ductiles anciennes, le long de

coulissements. Ces multiples consequences de

revolution de la Tethys de sa naissance a sa fermeture sur

les plates-formes cratoniques bordieres sont au cceur du

programme international « Peri-Tethys Programme ».

Une serie de memoires jalonncnt ces travaux et ceux

conduits depuis plusieurs ann£es. Le memoire Peri-Tethys

n° 1 etail consacre aux methodes d’etudes des plates-

formes. Les prochains seront consacres aux resultats

acquis sur ces plates-formes bordieres de la Tethys et des

chaines qui en sont issues.

Nous remercions le Museum national d’Histoire

naturelle pour avoir accepte d’inclure cette serie de

travaux dans sa collection renommee de monographies

scientifiques.

Ces memoires s’inscrivent dans l'ensemble des

etudes tethysiennes realisees par des equipes auxquclles

sont associes ou qu’animeront les coordonateurs et

participants de ces programmes. Les travaux antericurs de

references sont :

The chains around the Mediterranean and those

associated with them result from the changes that

occurred in the continental slope and Tethyan margins as

the ocean evolved. The formation of the flexural basins

in the foreland is a reaction of the cratons and micro¬

cratons of the Mediterranean lithospheric sill to crustal

thickening.

The effects of the pre-Tethyan distensions,

associated with accretion, were felt far away on the

bordering cratons well beyond the regions which were to

bend underneath the chains to become flexural basins.

Subsequently, in the collision phases between super -

cratons and the microcratons that were to constitute the

Mediterranean lithospheric sill, subsidence events

resumed along the ancient weak and ductile faults, along

strike-slip faults. These multiple repercussions of the

evolution of the Tethys, from its first opening to its

closure onto the bordering cratonic platforms, are the

crux of the International Peri-Tethys Programme.

A whole series of reports stand out as landmarks in

this research and work that has now been under way for

several years. The collection Peri-Tethys No I has been

devoted to methods of investigating the platforms.

Subsequent ones will bear on the results obtained on

these Tethyan margin platforms and the mountain ranges

that resulted from them.

We thank the Museum national d’Histoire naturelle

(Paris) for having included this series of works among its

collection of monographs of excellent composition.

These collections of work fit in well with the whole

body of research on the Tethys performed by research

teams of which the coordinators and participants of these

programmes are either members or leaders. Previous

works of reference are as follows:

ZlEGLER, P. A., 1990. — Geological Atlas of Western and Central Europe, Shell Internationale Petroleum Maatschappij,

distributed by Geological Society, London, Bath, 239 pp., 5 maps.

DERCOURT, J., RlCOU, L.-E. & Vrielynck, B. (eds), 1993. — Tethys Palaeogeographical Maps, Gauthier-Villars & Beicip,

Paris, distributed by Commission de la Carte Geologique du Monde, 77 rue Claude Bernard, Paris, 307 pp.,

14 maps.

ROURE, F. (ed), 1993. — Peri-Tethyan Platforms, Technip, Paris, 275 pp.

Nairn, A., RlCOU, L.-E. & Vrielynck, B. (eds), 1996. — The Ocean Basins and Margins, vol. 8, The Tethys Ocean

Plenum, New York & London, 530 pp.

Jean DERCOURT & Maunzio GAETANI

Source : MNHN, Paris

FOREWORD

The Peri-Tethys Memoir 2 developed out of the symposium on “Structure and Prospects of

Alpine Basins and Forelands", held during the American Association of Petroleum Geologists

International Conference and Exhibition October 17-20th. 1993 in The Hague, The Netherlands. This

two and one halfdays symposium was convocated by Peter A. Ziegler and Frank Horvath; it aimed

at discussing the structure, evolution and hydrocarbon habitat of the different segments of the

Alpine-Mediterranean fold-and-thrust belts and associated foreland and back-arc basins. The

program consisted of 33 papers covering the Alpine system from the Gibraltar Arc to the Black Sea

and the frontal thrust belt of the East Taurus. Abstracts of this symposium were published in the

American Association of Petroleum Geologists Bulletin, 77 (9), 1993. Participation of several

speakers from Eastern Europe and the former Soviet Union in this international symposium was

sponsored by Shell Internationale Petroleum Mij. B.V.

In the course of this symposium it was realized that papers presented provided a unique

insight into the data accumulated, mainly by the petroleum industry, during its exploration for oil

and gas in the different fold-and-thrust belts of the Alpine system and its associated basins. As this

information is vital to the understanding of the evolution of the Alpine system and the Peri-Tethyan

platforms, speakers were canvassed already during the symposium for a possible contribution of their

papers to the publication in a proceedings volume. In general the response was positive and often

immediate.

The next step was the task of the conveners to find an organization which was willing to

sponsor publication of such a volume. After the American Association of Petroleum Geologists had

declined sponsorship, the conveners were informed July 13th 1994 that the Executive Committee of

the international Peri-Tethys Program, headed by Prof. Dr. Jean Dercourt and Dr. Bernard Tissot,

had decided to sponsor publication of the symposium proceedings as Peri-Tethys Memoir 2, to be

edited by P. A. Ziegler and F. Horvath. With this, collection of manuscripts from contributors

commenced. Unfortunately, a number of papers presented during the symposium were withdrawn as

authors were assigned to different tasks in their organization or had changed to an other company.

Nevertheless, the editors were able to assemble 24 papers which provide a fairly comprehensive

coverage of the Alpine-Mediterranean system of fold-and-thrust belts and basins. This list includes

some papers which were not presented during the symposium but which were invited at a later stage

to cover areas of particular interest.

Papers included in this memoir come from the petroleum industry, from national research

agencies and from universities. These papers are in so far heterogeneous as the aspects emphasized

by the different authors vary considerably. However, all papers address the structural and

stratigraphic evolution of the respective area and its hydrocarbon habitat. All papers included in this

volume are based on recent compilations and integrations of surface — and sub-surface geological

and geophysical data, acquired in the context of hydrocarbon exploration and/or scientific research

programs aiming at understanding the architectures and origin of a specific orogen or basin. Much

of the data presented in this volume, particularly on East-European areas, was hitherto difficult to

access, partly due to its confidentiality and partly due to linguistic difficulties. In respect of the latter,

major and often very time consuming editorial ellorts were required.

The Peri-Tethys Memoir 2 is organized into an introductory and five regional chapters. In the

introductory chapter, the architecture of the Alpine orogen and its hydrocarbon systems are

14

FOREWORD

reviewed. In addition, a set of palaeogeographic maps is presented, retracing the Mesozoic opening

of Tethys and its Alpine closure, resulting in the suturing of Africa-Arabia and Europe. The

following chapters aim at providing the reader with a summary of the structural and stratigraphic

evolution, architecture and hydrocarbon potential of selected parts of the Alpine-Mediterranean

orogenic belt and its associated basins.

All papers were reviewed by the senior editor and most of them by one or more peer or outside

reviewers; thanks are extended to these reviewers for their efforts. The lay-out of this memoir was

taken care of by Dr Frank Horvath, Peter Szafian, Laszlo Lenkey and Orsolya Magyari at the

Lorand Eotvos University, Budapest. Editorial expenses were covered by the Peri-Tethys Program.

The editors would like to thank all contributing authors for their commitment to this project

and their sponsoring organizations for releasing the respective papers for publication. Special thanks

go to EXXON Exploration Company for absorbing the reproduction costs of the palaeogeographic

maps, given in the paper by P. O. Yilmaz et al., and to Elf-Aquitaine Exploration and Production

Company for financially supporting colour reproduction of text figures given in the paper by M. Le

Vot et al.

This volume could not have been produced without close cooperation between academic and

industrial Earth scientists. Such cooperation forms the basis of the on-going Peri-Tethys Program,

the sponsors of which are thanked for having agreed to the publication of this volume.

Peter A. Ziegler & Frank Horvath

Editors

List of reviewers

G. H. Bachmann

A. W. Bally

D. Bernoulli

G. Bertotti

G. Bessereau

S. Cloetingh

G. Gorin

B. Gunzenhauser

F. Horvath

P. Jordan

H. -P. Laubscher

P. E. R. Lovelock

A. Mascle

F. Roure

S. M. Schmid

M. Schwander

M. Stefanescu

M. A. Vul

P. A. Ziegler

Martin Luther University, Halle-Wittenberg, Germany

Rice University, Houston. Texas, USA

ETH. Zurich, Switzerland

Free University, Amsterdam, The Netherlands

Institut Frangais du Petrole. Rueil-Malmaison, France

Free University, Amsterdam, The Netherlands

University of Geneva, Geneva, Switzerland

PROSEIS AG, Zurich, Switzerland

Lorand Eotvos University, Budapest, Hungary

Amt fiir Wasserwirtschaft, Solothurn, Switzerland

University of Basel, Basel, Switzerland

Shell Internationale Petroleum Mij. BV, Den Haag, The Netherlands

Institut Frangais du Petrole, Rueil-Malmaison, France

Institut Frangais du Petrole, Rueil-Malmaison, France

University of Basel, Basel. Switzerland

Shell Internationale Petroleum Mij. BV, Den Haag. The Netherlands

Amoco Romania, Bucharest, Romania

UkrDGRI, Lvov, Ukraine

University of Basel, Basel, Switzerland

Architecture and petroleum systems of the Alpine orogen

and associated basins

P. A. Ziegler * & F. Roure**

* Geological-Paleontological Institute,

University of Basel, Bernoullistrasse 32,

CH-4056 Basel, Switzerland

** Institut Frangais du Petrole, 1-4 avenue de Bois-Preau,

BP 311, F-92506 Rueil-Malmaison Cedex, France

ABSTRACT

The Alpine orogen extends from Gibraltar to

the Black Sea and consists of an interlinking sys¬

tem of fold-and-thrust belts and associated foreland

and back-arc basins. Although these all evolved in

response to convergence of the Africa-Arabian and

European cratons, and coeval closure of the Tethys

oceanic basins, they differ widely in their architec¬

ture and evolutionary history in which such aspects

as orthogonal or oblique collision, the intensity of

collisional coupling of the evolving orogen with its

foreland, and the alternation of compression, trans-

pression and transtension, or even extension,

played a significant role.

Recently recorded deep seismic profiles image

the crustal architecture of some segments of the

Alpine orogen. Progressive, gentle orogenward

downflexing of the foreland crust, accommodating

foredeep basins, as well as localized crustal roots

are evident beneath the axial parts of the Pyrenees

and the Alps. However, in the northern and eastern

Carpathians, in Languedoc-Provence or in the

Betic Cordillera, the Moho remains horizontal or

even shallows toward the internal zones of these

orogens. This is thought to be related to syn-oro-

genic back-arc extension, as seen in the Pannonian

basin and the Tyrrhenian Sea, or to the opening of

the oceanic Algero-Provensal Basin. Seismic

tomography images deep lithospheric roots in dif¬

ferent segments of the Alpine orogen. Lithospheric

slabs appear to be still attached to the underthrust-

ed Apulian-Ionian, East-Mediterranean and

Moesian crusts beneath the southern Apennines-

Calabrian and Aegean arcs and along the south¬

eastern salient of the Carpathians, respectively.

Effective source rocks are widely distributed

within the Alpine orogen and in associated basins.

They are alternatively localized in pre-rift, syn-rift

or passive margin sequences, the age of which is

highly variable. In addition, syn-orogenic foreland

basin sequences can host important source-rocks

and are frequently the locus of biogenic gas gener¬

ation. Examples of syn-orogenic oil source rocks

are the Oligocene Menilite shales of the Carpathian

domain. The Po Plain and the Adriatic foreland

basin contain major biogenic gas reserves.

Contrasted thermal regimes and successive

episodes of sedimentary and tectonic burial

account for the great diversity of petroleum sys¬

tems identified in the Alpine orogen and associated

basins. Each hydrocarbon province is characterized

by a very distinct scenario for the timing of source-

ZlEGLER, P. A. & Roure, F\, 1996. Architecture and petroleum systems of the Alpine orogen and associated basins. In: Ziegler,

P. A. & Horvath, F. (eds), Peri-Tethys Memoir 2: Structure and Prospects of Alpine Basins and Forelands. Mem. Mus. natn. Hist. nat..

170: 15-45. Paris ISBN: 2-85653-507-0.

Source

16

P. A. ZIEGLER & F. ROURE: ALPINE OROGEN

rock maturation and petroleum expulsion, for

hydrocarbon migration from effective kitchens to

potential trapping domains, and for the preserva¬

tion of hydrocarbon accumulations.

INTRODUCTION

The Alpine orogen of the Mediterranean area

consists of a system of interlinking fold-and-thrust

belts and associated foreland and back-arc basins.

These differ in the age of their development and

deformation, their tectonic setting and architecture.

The evolution of all these features is intimately

linked with the Mesozoic break up of Permo-Trias-

sic Pangea, culminating in the step-wise opening of

oceanic basins forming the Western Tethys, and

their closure during the Alpine orogenic cycle.

For a long time, geologists have identified

within the Alpine allochthons remnants of the

Tethyan passive margins as well as ophiolites,

interpreted as remnants of coeval oceanic basins.

On the basis of these interpretations the former

plate boundaries between Europe, Apulia and

Africa were defined (Biju-Duval et al., 1977;

Bernoulli and Lemoine, 1980; Dercourt et al.,

1986; Favre and Stampfli, 1992). Oceanographic

surveys and off-shore drilling have shed light on

the origin of the various Mediterranean sub-basins.

These are variably floored by remnants of the for¬

mer Tethys Ocean (Ionian and Libyan seas), newly

formed oceanic crusts (Algero-Proven^al basin and

Tyrrhenian Sea), extended continental lithosphere

(Valencia trough, Aegean Sea) and thick continen¬

tal lithosphere (Adriatic Sea).

Available geological and geophysical data

sets, including recently recorded deep seismic pro¬

files, provide a 3D-image of the present crustal

architecture of the Alpine orogen. These data per¬

mit palinspastic restoration of former oceanic and

continental domains and. by applying plate tecton¬

ic concepts, a simulation of the plate kinematics

which underlay the evolution of the Alpine orogen.

Construction of regional palaeogeographic-palaeo-

tectonic maps has considerably advanced our

understanding of the evolution of the Alpine oro¬

gen and of the sedimentary basins which are asso¬

ciated with it. Yet, remaining uncertainties, for

instance about the configuration of the Tethys

embayment at the end of the Variscan orogeny

(how far to the East had collisional coupling

between Africa-Arabia and Europe progressed and

was the northeastern margin of Africa-Arabia

indeed fringed by an orogen), the timing of open¬

ing and closure of the various Mesozoic oceanic

basins forming the Western Tethys and the deriva¬

tion and dimension of some of the tectono-strati-

graphic units which are involved in the Alpine

orogen. account for major differences between

reconstructions proposed by, for instance, Ziegler

(1988, 1994a), Stampfli et al. (1991), Dercourt et

al. (1993) and Yilmaz et al. (this volume). It is the

objective of the on-going Peri-Tethys program to

resolve some of these outstanding questions.

As an introduction to this volume, in which

selected case history studies on basins and fold-

and-thrust belts and petroleum provinces are dis¬

cussed, this paper aims at providing an overview of

the crustal architecture and evolution of the entire

Alpine orogen and its associated basins. The role

played by Tethyan syn- and post-rift and Alpine

syn-orogenic series in the development of petrole¬

um systems will be discussed with reference to the

regional examples documented in the following

chapters.

ALPINE OROGENS AND ASSOCIATED

BASINS

The complex, arcuate geometry of the Alpine

orogen was preconditioned by the rift-induced con¬

figuration of its forelands, the pattern of oceanic

basins and intervening microcontinental blocks and

the kinematics of their interaction during the

Alpine convergence of Africa-Arabia and Eurasia.

The major elements of the Alpine orogen evolved

by closure of oceanic basins of variable size and

age; such sutures are characterized by internal

ophiolitic zones and major nappes. However, a

number of fold belts developed by inversion of

intra-continental rift zones (e.g. Atlas, Celtiberian

Source :

PERI-TETHYS MEMOIR 2: ALPINE BASINS AND FORELANDS

17

range, Provence-Languedoc. Dauphine, Dobrogea,

Crimea). A transitional feature between these two

end members are the Pyrenees which evolved out

of an inter-continental transform rift zone, charac¬

terized by localized mantle denudation.

Remnants of Mesozoic Tethys Versus Neogene

Oceanic Crust

According to geophysical data and palinspas-

tic reconstructions, remnants of the Mesozoic

Tethys Ocean are still preserved in the Ionian Sea

and probably also in the northern parts of the East¬

ern Mediterranean. Both domains are at present

still being subducted beneath the Calabrian and the

Cretan and Cyprus arcs, respectively (Fig. 1).

Thick Mesozoic and Cenozoic sediments prohibit

direct sampling and dating of the oceanic crust in

these areas, which could be as old as Permian (first

deep marine sediments on the Pelagian block and

pelagic series of the Hawasina nappes in Oman;

Stampfli et al., 1991; Stampfli, 1996), or as young

as Cretaceous (i.e. coeval with the onset of the

rotation of Apulia with respect to Africa and rifting

of the Syrte Basin in Libya). In the Eastern

Mediterranean, the area occupied by oceanic crust

is uncertain (Makris et al., 1983; Sage and

Letouzey, 1990) and its age is still debated; similar

to most of the ophiolitic units accreted in the

Alpine orogen, the oceanic part of the East-

Mediterranean basin is probably of mainly Triassic

to Jurassic age. However, lack of hard information

on the nature and age of the East-Mediterranean

crust provides for major uncertainties and differ¬

ences in palaeo-reconstructions retracing the open¬

ing of the Western Tethys and its closure during the

Alpine orogenic cycle (Ziegler, 1988; Stampfli,

1996; Dercourt et al., 1993; Yilmaz et al, this vol¬

ume).

In contrast, the oceanic crust of the Western

Mediterranean is Neogene in age. Opening of the

oceanic Algero-Provengal Basin is dated as Burdi-

galian by the transition from syn-rift to thermal

post-rift subsidence of its margins and by palaeo-

magnetic data (Vially and Tremolieres, this vol¬

ume). Oligocene-earliest Miocene rifting in the

domain of the Gulf of Lyons and the Valencia

Trough, culminating in opening of the Provencal

Basin, was contemporaneous with northwest-dip¬

ping subduction of the Alboran-Ligurian-Piemont

Ocean and thus initiated in a back-arc setting

(Maillard and Mauffret, 1993). However, sea-floor

spreading in the Algero-Provengal Basin cannot be

related to back-arc extension (Ziegler, 1994b). On

the other hand. Late Miocene and younger rifting

in eastern Sardinia and in Tuscany (Keller et al.,

1994; Spadini et al., 1995), as well as magnetic

anomalies and results of deep-sea drilling, date

opening of limited oceanic domains in the Tyrrhen¬

ian Sea as Pliocene and Quaternary (Wezel, 1985).

This young oceanic basin developed in a back-arc

setting with respects to the Apennine orogen; rift¬

ing and opening of the Tyrrhenian Basin behind

and above the west-dipping Apennine subduction

zone was contemporaneous with continued sub¬

duction of the Apulian-Ionian crust (Serri et al.,

1993).

Back-arc extension, even if it did not always

culminate in opening of oceanic basins, as e.g. in

the Aegean Sea (Jolivet et al., 1994) and the Pan-

nonian Basins (Royden and Horvath, 1988; Tari et

al., 1992), causes rapid subsidence of formerly ele¬

vated compressional structures (negative inver¬

sion). Under such conditions, pre-existing

compressional detachment faults can be tensionally

reactivated, as seen in the Oligocene-early

Miocene evolution of the the Languedoc coast

(Vially and Tremolieres, this volume) and in the

Miocene development of the Danube Basin (Tari,

this volume).

Ophiolitic Sutures

Unlike the West- and East-Mediterranean

basins, which both still comprise undeformed

oceanic crust, the Alpine orogen is characterized

by several systems of ophiolitic bodies; these rep¬

resent tectonized and obducted remnants of former

Mesozoic oceanic basins, presently localized in

often narrow suture zones within the allochthon.

As outcropping ophiolitic nappes represent only

fragments of these oceanic basins, they do not nec-

essarely record the onset of sea- floor spreading in

the respective basins, the timing of which must be

FIG. 1. Presenl distribution of oceanic sutures and relict oceanic domains within the Alpine orogen.

18

P. A. ZIEGLER & F. ROURE: ALPINE OROGEN

Source : MNHN. Paris

PERI-TETHYS MEMOIR 2: ALPINE BASINS AND FORELANDS

19

derived from the sedimentary record of the offset¬

ting passive margin prisms, now involved in the

Alpine nappes. Nevertheless, these ophiolitic

suture zones relate to distinct segments of the for¬

mer Tethys and record its step-wise opening

(Fig. 1).

During the Late Permian-Early Jurassic initial

break-up phase of Pangea, the Hallstatt-Meliata-

Mures, Vardar and Sub-Pelagonian and possibly

also the East-Mediterranean oceanic basins

opened. To the east, this system of oceanic basins

finds its continuation in the Izmir-Ankara-Erzincan

and the Taurides ophiolitic belts. Opening of this

system of oceanic basins resulted in partial separa¬

tion of the Italo-Dinarides-Anatolia bock from

Europe. With the Middle Jurassic development of a

discrete transform/divergent plate boundary

between Gondwana and Laurentia, opening of the

Central Atlantic entailed a change in the opening

kinematics of the Western Tethys; these were dom¬

inated during the Late Jurassic-Early Cretaceous

by a major sinistral translation between Africa-

Arabia and Europe, inducing progressive

transtensional opening of the Alboran-Ligurian-

Piemont-South Penninic Ocean. This was accom¬

panied by decoupling of the Apulian terrane from

Africa-Arabia and its complete isolation . At the

same time new subduction systems developed in

the Vardar-Hallstatt system of oceanic basins, gov¬

erning their gradual closure. Early Cretaceous

opening of the North Atlantic was paired with

opening of the Bay of Biscay and of the North Pen¬

ninic Valais Trough, which may equate to the intra-

Carpathian Magura-Piennidic zone (Ziegler, 1988;

Dercourt et al., 1993).

Some elements of the Alpine orogenic system,

such as the Pyrenees and the Greater Caucasus,

lack true ophiolitic sutures, despite the fact that

they developed by closure of pre-existing smaller

or larger oceanic troughs. In such fold belts ultra-

mafic rocks may occur locally as narrow tectonic

slices.

Closure of the different segments of the

oceanic Tethys basins during the Alpine orogenic

cycle was diachronous, also along the trace of the

individual basins For example, in the Central Alps,

the South Penninic Ocean was closed during the

early Paleocene whereas in the Western Alps, linal

closure of the Piemont-Ligurian Ocean occurred

only during the late Eocene (Ziegler et al., this vol¬

ume).

Foreland Flexure and Residual Crustal Roots

Deep seismic profiling, including reflection

and wide-angle surveys, gravity data, seismic

tomography and modelling have provided a better

insight into the crustal and lithospheric architecture

of the various segments of the Alpine orogen.

Palaeomagnetic data, the inventory of sea¬

floor magnetic anomalies and palinspastic recon¬

structions of the different segments of the Alpine

orogen indicate that Late Mesozoic and Paleogene

convergence of Africa-Arabia and Europe amount¬

ed to hundreds of kilometres and that it was

accommodated by the subduction of equivalent

amounts of oceanic and partly also continental

lithospheric material (de Jong et al., 1993). This

concept is supported by presence of long lithos¬

pheric slabs, penetrating the asthenosphere, which

are imaged by the distribution of earthquakes and

by seismic tomography, for instance beneath the

active Calabrian and Aegean arcs as well as in the

western Mediterranean (Spakman, 1990; Wortel et

al., 1990; Spakman et al., 1993). Alternatively,

major slab detachments probably occurred beneath

the Alps, the Dinarides and the Carpathians during

or soon after the Cretaceous-Paleogene episodes of

intense shortening (von Blankenburg and Davies,

1995) .

In contrast to major subduction zones, the

Pyrenees are characterized by a limited crustal root

only; this is in agreement with the relatively small

lithospheric contraction during the late Senonian-

Paleogene Pyrenean orogeny (about 1 10 km). Sim¬

ilarly, about 120 km of Oligocene to Recent

subduction of the European continental lithosphere

beneath the Alps, accounts for their present crustal

root and a corresponding new subduction slab (de

Jong et al., 1993; ECORS Pyrenees team, 1988;

Choukroune et al., 1989; Frei et al., 1989; Roure et

al., 1989, 1990; Pfiffner et al., 1988; Schmid et al.,

1996) . In both cases, conjugate foreland basins

developed on either side of the orogen, although

seismic and gravimetric data attest for strong

asymmetries at depth (Bayer et al., 1989). Whereas

20

P A. ZIEGLER & F. ROURE: ALPINE OROGEN

in the Pyrenees the Iberian infra-continental mantle

was progressively subducted northward, the Euro¬

pean lithosphere was underthrusted to the southeast

and south beneath the Western and Central Alps.

Other segments of the Alpine chain, such as

the Languedoc-Provence, the Western and Eastern

Carpathians and also parts of the Apennines, are

characterized by limited crustal roots and a Moho

which progressively shallows toward the internal

zone of these orogens, corresponding to the Gulf of

Lions, the Pannonian Basin and the Tyrrhenian

Sea, respectively (Figs. 2 and 3; Tomek, 1993;

Szafian et al., 1995). Frequently interpreted as a

new Moho, this geometry of the crust-mantle

boundary probably results from the extensional

collapse of the internal parts of these belts, involv¬

ing extensional reactivation of pre-existing com-

pressional detachment horizons (thrust faults) and

progressive denudation of the lower crust (meta-

morphic core complexes). Under post-orogenic

conditions, this can entail rapid uplift and erosion

of the external parts of the orogen and unflexing of

the foreland lithosphere, as seen, for instance, in

the northern Carpathians.

Foredeep basins flanking the Apennines and

the southeastern Carpathians are extremely deep

and contain up to 10 km of Miocene to Pliocene

sediments. In contrast, the eastern Pyrenean and

the West-Alpine forelands are almost devoid of

syn-flexural sediments.

The width and depth of a flexural basin large¬

ly depends on the thickness and thermal regime of

the foreland lithosphere, controlling its rheology,

as well as on the loads which are exerted on it by

the subduction slab and the overriding orogenic

wedge (Watts et al., 1982; Kusznir and Karner,

1985; Desegaulx et al., 1991; Doglioni, 1993).

Flexural down-bending of thick continental lithos¬

phere can be accompanied by the development of

an array of relatively small tensional, essentially

basin-parallel normal faults at upper crustal levels,

as evident in the German and Austrian parts of the

Molasse Basin (Roeder and Bachmann, Zimmer

and Wessely, this volume). If such a foreland crust

is weakened by pre-existing faults which can be

reactivated during its flexural deformation, strain

may be concentrated on a few major faults which

can have throws of the order of 1 km and more, as

seen in the Ukrainian part of the Carpathian fore-

deep (Sovchik and Vul, this volume). Rapidly and

deeply subsiding foredeep basins are generally

characterized by strongly stretched and attenuated

continental crust which is thinner and warmer than

adjacent parts of the foreland. This accounts for the

contrasted geometries of the western and eastern

portions of the Aquitaine foredeep, and for the seg¬

mentation of the Periadriatic depressions in front

of the Apennines or between the Dinarides and the

Albanides (Royden et al., 1987; Desegaulx et al.,

1991; Doglioni, 1993).

In some foreland basins, intra-plate congres¬

sional structures play an important role and, by

their development, can either impede the develop¬

ment of a flexural basin or can cause partial or total

destruction of a pre-existing flexural basin. Such

structures can develop by inversion of pre-existing

tensional basins, in which case they may involve

the basement (i.e. Provence, Dauphine; Roure and

Colletta, this volume). Alternatively, activation of

sedimentary detachment horizons within the fore¬

land basin can cause the development of thin-

skinned compressional structures, as seen in the

Prerifaine chains of Morocco (Zizi, this volume),

the South-Alpine Lombardian Basin (Ziegler et al.,

this volume), the Jura Mountains (Philippe et al,

this volume) and also in the Apennine foredeep

(Anelli et al., this volume).

Oblique Versus Orthogonal Collision and Strain

Partitioning

Although Africa-Arabia converged with Eura¬

sia since Senonian times in a north-south directed

counter-clockwise rotational mode, sinistral

motions between them, related to the opening of

the Atlantic Ocean, decreased gradually and ceased

at the transition from the Paleocene to the Eocene

in conjunction with opening of the Norwegian-

Greenland Sea. However, during the Oligocene and

Miocene a dextral component is evident in their

convergence pattern; this translatory movement

probably persisted into Recent times, as indicated

by earthquake focal mechanisms and neotectonic

deformations (Ziegler, 1988, 1990).

During the Late Jurassic-Early Cretaceous

opening phases of the Central Atlantic, sinistral

motion between Africa-Arabia and Europe gov-

Source :

PERI-TETHYS MEMOIR 2: ALPINE BASINS AND FORELANDS

21

Source : MNHN, Paris

FIG. 2. Block diagram outlining ihe role of strain-partitioning in areas of oblique convergence or collision

22

P. A. ZIEGLER & F. ROURE: ALPINE OROGEN

erned the closure of the Hallstatt-Vardar ocean sys¬

tem, rotation of the Italo-Dinarides-Anatolia block

and its progressive incorporation into the Dinar-

ides-Hellenides-Pontides orogenic system. With

the Senonian onset of northward drift of Africa-

Arabia and progressive opening of the Atlantic,

these sinistral motions gradually decreased where¬

as increasing space constraints in the Western

Tethys caused rapid westward propagation of sub-

duction zones into the the domain of the Ligurian-

Alboran Ocean. For instance, in the Alps, the

Cretaceous orogenic cycle was characterized by

northwest-directed mass transport whereas the

Paleogene orogeny was dominated by north direct¬

ed mass transport (Schmid et ah, 1996). Paleogene

progressive closure of the Tethys and increasing

collisional coupling of the evolving orogen with its

forelands, lateral block escapes and oblique

motions played an increasingly important role.

Suturing of Iberia to Europe was accompanied by

the development of the left-lateral North Pyrenean

Fault. Eastward directed Oligo-Miocene mass

transport from the Alpine into the Carpathian

domain, as a consequence of full-scale collision of

the Adriatic indenter with Europe (Ratschbacher et

al., 1991), was accompanied by left-lateral motions

along the North Carpathian Pienniny Klippen belt

and incipient right-lateral motion along the South

Carpathian foothills (Fig. 1; Laubscher, 1992a;

Ellouz and Roca, 1994). Miocene-Pliocene devel¬

opment of a system of right-lateral strike slip fault

systems, including the South Atlas fracture system,

the Insubric line of the Southern Alps, the intra-

Dinarides Peri-Adriatic and the North Anatolian

fault systems, may be partly related to the dextral

translation of Africa-Arabia and Europe and partly

to lateral mass redistribution in response to the

massive indentation of Arabian and the Adriatic

block (Ziegler, 1988).

However, as in still presently active transpres-

sional orogens, such as the South Caribbean belt in

Eastern Venezuela, or along the San Andreas Fault

west of the San Joachin Valley, a strong strain par¬

titioning occurred within most of the Alpine oro¬

genic belts (Laubscher, 1992a; Passalacqua et al.,

1995). Thus, the overall northwest- or northeast¬

trending collision zone between major plates was

ultimately confined to the subducted lithosphere in

the footwall, and to the hanging-wall mantel inden¬

ter (Fig. 2). In contrast, frontal accretion character¬

ized the conjugate external thrust belts on both

sides of the orogenic wedge, with major thrusts

always paralleling the active plate boundary (see

e.g. Gilmour and Makel, Le Vot et al., Philippe et

al, this volume). Apparently, the oblique conver¬

gence component was largely absorbed within the

orogenic wedge by strike-slip motions along, at

shallow levels, sub-vertical faults. Such faults can

strike normal to the overall convergence direction

as e.g. the Insubric line of the Central Alps, or at a

high-angle as e.g. the intra-Dinarides Peri-Adriatic

line. Their orientation is largely controlled by the

geometry of rigid indenters, such as the Apulian

platform.

o

10

20

30

40

50

Duplexes

made up of

Iberian lower

crust

S Ebro

Basin

North Pyrenean

fault

N PF

Aquitaine N

Basin

0

10

20

30

40

50km

FIG. 3. Crustal section across the Pyrenees

Source :

PERI-TETHYS MEMOIR 2: ALPINE BASINS AND FORELANDS

23

RIFTING AND DEVELOPMENT OF

PASSIVE MARGINS

Tethyan rifting initiated soon after the consoli¬

dation of the Variscan orogen at the end of the

Westphalian (von Raumer and Neubauer, 1993).

Development of the widespread, wrench- and rift-

induced Permo-Carboniferous successor basins of

Western and Central Europe Europe can be related

to such processes as changes in the convergence

pattern of Gondwana and Laurussia during the

Allegenian phase of the Appalachian-Mauretanides

orogen, to back-arc extension in response to roll¬

back of the Variscan subduction slab and its ulti¬

mate detachment from the lithosphere and to the

onset of regional extension, controlling the break¬

up of Pangea (Ziegler, 1990, 1993).

In the Western Tethys domain, major crustal

extension commenced, however, only after the

Appalachian-Variscan suture between Gondwana

and Laurussia had become inactive at the transition

to the Late Permian (Ziegler, 1989; Stampfli and

Pillevuit, 1993). Rifting activity accelerated during

the Triassic, propagated westward and interfered in

the Atlantic domain with the southward propagat¬

ing Arctic-North Atlantic rift system (Ziegler,

1988). Step-wise opening of the different oceanic

basins of the Western Tethys, resulting in the

development of passive margins, as summarized

above, is retraced in the palaeogeographic/palaeo-

tectonic maps of Yilmaz et al. (this volume). We

recall here only, that probably most of the Tethyan

passive margins developed during the Triassic and

the Middle Jurassic.

Variscan Inheritance and Localization of Tethys

Rifts

The Triassic-Jurassic Gulf of Mexico-Central

Atlantic-Western Tethys rift/transform system rep¬

resents one of the major break-up axes of Permo-

Triassic Pangea. Significantly, this break-up axis

coincides to a large extent with the Appalachian-

Variscan suture of Gondwana and Laurussia

(Ziegler, 1990, 1993). Rheological considerations

suggest that the orogenically destabilized lithos¬

phere of this suture was considerably weaker than

that of the flanking cratons (Cloetingh and Banda,

1992) and, as such, preconditioned the localization

of this break-up axis.

In Western and Central Europe and the West¬

ern Tethys domain, the lithosphere was further

weakened during the Stephanian-Autunian collapse

of the orogenically thickened Variscan crust.

Wrench- and extensional faulting, resulting in the

uplift of core-complexes, was associated with the

synkinematic intrusion of granites and the extru¬

sion of alkaline and calc-alkaline magmas. This

was presumably accompanied by the detachment

of the subducted lithospheric slab(s), at least partial

delamination of the deep lithospheric roots of the

Variscan orogen and its corresponding uplift.

Moreover, transtensional uplift of core-complexes

and erosional unroofing of the crust was paralleled

by upwelling of the asthenosphere and the interac¬

tion of mafic melts with the crust. With the termi¬

nation of wrench activity at the transition to the

Late Permian, cooling of the thermally destabilized

lithosphere controlled crustal subsidence and

regional transgressions. In time, the crust/mantle

boundary re-equilibrated regionally at depths of

about 35-40 km (Ziegler, 1990; Ziegler et al.,

1995; Costa and Rey, 1995).

In the European Alpine foreland there is

ample evidence of repeated Mesozoic reactivation

of Permo-Carboniferous crustal discontinuities, in

part guiding the localization of rifted structures,

such as the Polish Trough (Ziegler, 1990). Deep

reflection-seismic profiles show that Variscan com-

pressional structures were only rarely reactivated

during the Mesozoic rifling phases in the distal

parts of the Alpine forelands.

For instance, the ECORS seismic profiles

image beneath the conjugate Ebro and Aquitaine

forelands of the Pyrenees south-verging Variscan

structures (Scrvey Geologic de Catalunya, 1993).

These were transected by Permo-Carboniferous

wrench faults and apparently did not materially

contribute towards the localization of Mesozoic

extensional faults. In contrast, reactivation of

south-verging Hercynian structures probably

accounts for coaxial Alpine deformations along the

southern flank of the Pyrenees (Desegaulx et al.,

1990). On the other hand, the Mesozoic Bay of

24

P. A. ZIEGLER & F. ROURE: ALPINE OROGEN

Biscay rift zone appears to be superimposed on a

major Permo-Carboniferous wrench zone.

Although the architecture of those parts of the

Variscan orogen which were overprinted by Alpine

deformations is still poorly known (von Raumer

and Neubauer, 1993), it is evident that also these

areas were affected by intense Permo-Carbonifer¬

ous wrench tectonics and associated magmatism. It

is likely that also in these areas reactivation of

Permo-Carboniferous crustal discontinuities played

a significant role in the localization of the Tethys

rift systems.

Late Carboniferous coal-measures contained

in Variscan foreland and successor basins, as well

as coal-measures and lacustrine shale of the

Stephanian-Autunian wrench-induced troughs, pro¬

vide potential source-rocks in the pre-rift sedimen¬

tary sequences of the European foreland.

Depending on their burial beneath the Tethyan pas¬

sive margin sequence, or beneath syn-orogenic

flexural sequences and the Alpine allochthon, such

source-rocks can have locally preserved part of

their petroleum potential and thus can contribute to

effective petroleum systems, as for instance in the

Jura Mountains and in the subthrust play of the

Polish Carpathians (Bessereau et a!., this volume).

Rifting and Development of Passive Margins

As discussed above, rifting activity in the

Tethys domain spanned Permian to Early Creta¬

ceous times and culminated in the step-wise open¬

ing of its constituent oceanic basins. Late Permian,

Triassic and Early Jurassic rifting activity is well

documented in the different parts of the Western

Tethys (Stampfli and Pillevuit, 1993; Stampfli,

1996). During the Triassic, rifting activity propa¬

gated westwards and affected very wide areas

around the future zones of crustal separation

(Ziegler, 1988). In time, rifting activity concentrat¬

ed on zones of future crustal separation, as seen for

instance in the Southern Alps (Bertotti et al.,

1993). From Mid-Jurassic to Early Cretaceous

times, rift and wrench activity in the Western

Tethys was governed by the sinistral translation of

Africa-Arabia relative to Europe, in response to

progressive opening of the Atlantic Ocean.

Earliest passive margins were associated with

the opening of the Hallstatt-Vardar system of

oceanic basins. Mid-Jurassic opening of the Albo-

ran-Ligurian-Piemont-South Penninic ocean result¬

ed in the development of a new set of passive

margins. Late Jurassic-Early Cretaceous opening

of the Bay of Biscay, the North Penninic and the

Magura basins led to the development of yet an

other system of passive margins.

Development of the different Tethys rift sys¬

tems entailed a renewed destabilization of the

asthenosphere-lithosphere system. Crustal exten¬

sion was accompanied by variable levels of rift

magmatism. As there is no obvious evidence for

hot-spot activity, crustal extension was presumably

of a “passive” nature, driven by far field stresses

governing the break-up of Pangea (Ziegler, 1993,

1995a). The availability of pre-existing crustal dis¬

continuities, which could be tensionally reactivat¬

ed, favoured simple shear crustal extension and the

development of upper and lower plate margins

(Favre and Stampfli, 1992). Locally crustal stretch¬

ing involved the activation of intra-sedimentary

detachment levels; for instance, in the basin of

Southeastern France, low-angle listric extensional

faults, soling out in Carboniferous coal-measures

or Triassic evaporites, account for decoupling of

the sedimentary cover structures from the base¬

ment. Subsidence of grabens and half-grabens was

accompanied by uplift of the rift shoulders and the

development of rift-flank basins (Favre and

Stampfli, 1992). Upon achievement of crustal sep¬

aration, marginal graben systems became inactive

and were incorporated into the newly formed pas¬

sive margins.

The duration of the rifting stage of the differ¬

ent Tethys extensional systems is highly variable.

For example, the Aquitaine-Bay of Biscay basin

records some 130 Ma of intermittent rifting activi¬

ty prior to its Mid-Aptian transtensional opening

(Desegaulx and Brunet, 1990), whereas the Atlas

troughs, after about 60 Ma of rifting activity,

became inactive in conjunction with crustal separa¬

tion in the Central Atlantic.

During much of Late Permian to Early Creta¬

ceous times, large parts of the Tethys domain were

dominated by carbonate shelves. Syn-sedimentary

tectonics accounted for rapid lateral facies and

thickness changes. Partly reefal or dolomitic shal¬

low water carbonates were restricted to the rift

Source :

PERI-TETHYS MEMOIR 2: ALPINE BASINS AND FORELANDS

25

flanks, the crests of tilted blocks and to little

extended platforms, whereas shaly or cherty deeper

water carbonates and shales were deposited, partly

under anoxic conditions, in rapidly subsiding

grabens and in sediment starved lagoons.

Triassic, Jurassic and Early Cretaceous syn-

rift source rocks provide for an number of effective

petroleum systems (Table I ).

How Passive Were the Tethys Margins

The present day Central Atlantic margins

remained in a passive setting for 180 Ma. Also the

Arabian Shelf had undergone a passive margin

evolution for some 175 Ma before it became colli-

sionally coupled during the Senonian with the

evolving Zagros orogen. In contrast the passive

margins of the Western Tethys had a relatively

short life span. For instance, the eastern margin of

the Italo-Dinarides block was incorporated into the

Dinarides orogen about 100 Ma after the Vardar

Ocean had opened (Frasheri et al., this volume).

The Austroalpine margin, facing the South Pen-

ninic Ocean, remained in a passive setting for

some 60 Ma before it was converted into an active

margin during the Early Cretaceous. Intra-Senon-

ian and Paleocene compressional deformation of

the East-Alpine and North-Carpathian forelands

occurred about 50 Ma after the Valais and Magura

basins had opened (Kovac et al., 1993; Ziegler et

al., this volume). On the other hand, the Pyrenean

margin was only for some 25 Ma transtensionally

“passive" before it was compressionally deformed

(Le Vot et al., this volume).

In view of the opening kinematics of the

Atlantic Ocean and of the Tethyan system of

oceanic basins, controlling also the interaction of

the different blocks delimited by the latter, many of

the evolving Tethyan passive margins were repeat¬

edly tectonically destabilized. This had repercus¬

sions on their thermal subsidence pattern and the

resulting development of passive margin sedimen¬

tary prisms. Examples of tectonic destabilization of

Tethyan passive margins are:

( 1 ) Late Jurassic and earliest Cretaceous

wrench activity in the Bohemian massif

and southward adjacent areas in conjunc¬

tion with a stress reorganization in the

North Sea rift system (Ziegler, 1990)

TABLE 1

Effective source rocks in Alpine Orogen and associated basins

26

P. A. ZIEGLER & F. ROURE: ALPINE OROGEN

(2) Early Cretaceous opening of the Valais

Trough, disrupting the subsidence pattern

of the European shelf bordering the South-

Penninic trough which had begun to open

during Mid-Jurassic times (Stampfli, 1993)

(3) Late Jurassic-Early Cretaceous wrench-

induced deformation of the Moroccan

shelf during the opening of the Central

Atlantic and Alboran-South Penninic

oceans (Favre and Stampfli, 1992)

(4) Early Cretaceous rifting in Libya and

Egypt. It is likely that also thermal subsi¬

dence of the Italo-Dinarides block was

influenced by its Early Cretaceous rota¬

tion.

The timing of incorporation of the different

peri- and intra-Tethyan passive margin prisms into

syn-orogenic flexural basins is also highly vari¬

able. Whereas during Cretaceous times progres¬

sively more internal units of the Italo-Dinarides

Block were incorporated into the Hellenic-Dinar-

ides foreland basin, the western shelves of this

block were incorporated into the Apennine fore¬

deep only during the Late Oligocene and Miocene

(Frasheri et al., Anelli et al., this volume). A spe¬

cial case is presented by the Helvetic shelf of the

Central and Eastern Alps which began to subside

during the Eocene under the load of the advancing

nappes; however, pre- or even syn-collisional

stresses exerted on this shelf induced transpres-

sional reactivation of pre-existing crustal disconti¬

nuities and profound disruption of its sedimentary

cover. For instance, intra-Senonian and Paleocene

inversion of the Polish Trough and the Bohemian

Massif resulted in partial destruction of the passive

margin prism of the East-Alpine and the Northl¬

and East-Carpathian forelands (Ziegler, 1990;

Ziegler et al., 1995).

Correspondingly, the age, thickness and com¬

position of the different peri- and intra-Tethyan

passive margin prisms is highly variable. These

passive margin prisms attained maximum thickness

prior to their incorporation into flexural basins.

Correspondingly, potential source-rocks of the syn-

rift and passive margin sequences reached maxi¬

mum sedimentary burial and started to generate

hydrocarbons already prior to their subsequent bur¬

ial beneath flexural foredeep prisms or the Alpine

allochthonous units. In distal foreland areas, which

escaped such burial, the syn-rift and passive mar¬

gin sedimentary series preserved part of their

petroleum potential (e.g. Aquitaine Basin, Le Vot,

this volume).

The Tethyan passive margin sequences

account for the source-rocks, reservoirs and/or

seals of some of the most efficient petroleum sys¬

tems identified within the external parts of the

Alpine fold-and-thrust belts and their forelands

(Table 1). Examples are the Early Jurassic Posido-

nia shales of the Albanian Ionian zone (Baudin and

Lachkar, 1990; Frasheri et al., this volume) and the

Aquitaine basin (Le Vot et al., this volume) and the

Late Jurassic shales beneath the Vienna basin

(Ladwein et al., 1991). Although numerous Creta¬

ceous black-shale intervals have been reported

from the peri-Tethyan passive margins, particularly

from the Albian-Cenomanian and Turonian series

of the Periadriatic domain, in Aquitaine or in the

Carpathians (Le Vot et al., Stefanescu and Baltes,

this volume), their efficiency is still questionable

as only limited hydrocarbon accumulations could

be directly correlated with them.

CONVERGENCE, COLLISION AND

SUBDUCTION OF CONTINENTAL

LITHOSPHERE

We recapitulate that in the Western Tethys ini¬

tiation of and activity along subduction zones was

controlled during the Late Jurassic-Cretaceous

Pangea break-up phase by the sinistral translation

of Africa-Arabia and Europe, and during Senonian

to Recent times by their Alpine convergence. As

the European and Africa-Arabian passive Tethys

margins were not parallel, and as the Iberian and

the Italo-Dinarides microcontinents acted as inden¬

ted during the Alpine cycle, collisional events

were diachronous (Tapponnier, 1977). Preservation

of remnants of Tethyan oceanic crust in the Ionian

Sea and in the Eastern Mediterranean illustrates

that continent-to-continent collision has not yet

occurred along the entire trace of the Tethys suture

(Fig. 1).

Source :

PERI-TETHYS MEMOIR 2: ALPINE BASINS AND FORELANDS

27

The onset of subduction processes can be

stratigraphically dated by the sedimentary record

of accretion prisms, or petrologically and radio-

metrically by arc volcanism and HP/LT metamor¬

phism. The collision of an arc trench system with a

passive margin, marking the transition from sub¬

duction of oceanic to continental lithosphere (B-

type to A-type subduction), is reflected by the

rapid subsidence of the foreland crust and by the

deposition of syn-orogenic flysch series on top of

the passive margin sequence. However, it must be

realized, that initial collision zones are generally

deeply buried beneath orogenic wedges and in part

have been subducted to great depths. Only in rare

cases will subduction progradation, or perhaps

extension, entail the exhumation of an earlier

active subduction zone and render it accessible to

surface observation (e.g. Tauern window;

Froitzheim et al., 1996). These criteria have been

applied in the construction of palaeo-

geographic/palaeotectonic maps retracing the evo¬

lution of the Alpine orogen (Ziegler, 1988;

Dercourt et al., 1993; Stampfli, 1996; Yilmaz et al,

this volume).

Not only the timing of orogenic activity is

highly variable in the different segments of the

Alpine system but also the amount of crustal short¬

ening, including subduction of oceanic crust and

post-collisional subduction of continental lithos¬

phere. In this respect, the intensity of collisional

coupling of the evolving orogenic wedge with its

foreland, as well as the availability of a thick sedi¬

mentary cover which could be detached from the

foreland crust, played an important role in the

architecture of the different fold-and-thrust belts

forming the Mediterranean-Alpine orogen. In the

following selected examples are discussed which

are further analyzed in the different chapters of this

volume.

Pyrenees

The Pyrenees evolved in response to late

Senonian-Paleogene transpressional closure ot the

wedge-shaped inter-continental Bay of Biscay

Basin; this was induced by the build-up of far-field

compressional stresses related to the convergence

of Africa with Europe, causing clock-wise rotation

and escape of Iberia. Crustal shortening across the

Pyrenees amounts to some 110 km (Roure et al.,

1989; Desegaulx et al., 1990).

Their axial zone is characterized by asymmet¬

rically north- and south-verging, thrusted basement

blocks. Their northern foreland, the Aquitaine

Basin, is characterized by partly reactivated exten-

sional fault bocks and thin skinned thrust sheets

which are detached from the basement at the level

of Triassic evaporites (Le Vot et al., this volume).

The southern Pyrenean external zone is character¬

ized by thin skinned thrust sheets, detached from

their basement at the level of Triassic evaporites;

the basement cores of these sheets form part of the

internal structure of the Pyrenees (Verges and

Munoz, 1990). The Pyrenean orogeny is coeval

with orogenic activity in the Betic Cordillera. Pale¬

ogene compressional stresses exerted on cratonic

Iberia caused reactivation of Permo-Carboniferous

and Mesozoic crustal discontinuities, controlling

inversion of the Celt-Iberian and Catalonian Coast

ranges and upthrusting of the Sierra Guadarrama

basement block (Ziegler, 1988; Salas and Casas,

1993; Ziegler et al., 1995).

Commensurate with the amount of crustal

shortening achieved across the Pyrenees, they are

characterized by a limited north-verging crustal

root and no pronounced lithospheric root; they lack

a syn-orogenic magmatism (Fig. 3; Spakman,

1990; Servey Geologic de Catalunya, 1993).

Whereas the northwestern, deep foreland

basin of the Pyrenees hosts the major Aquitaine

hydrocarbon province (Le Vot et al., this volume),

its eastern shallower parts and the southern fore¬

land contain no significant hydrocarbon accumula¬

tions.

Western Alps

From plate reconstructions, the total crustal

shortening achieved across the Western Alps is

estimated to amount to some 400 km (Platt et al.,

1989). This involved Late Cretaceous activation of

the Apulian margin. Late Cretaceous-Paleogene

closure of the oceanic Piemont, Eocene closure of

the Valais basins (Froitzheim et al., 1996) and sub-

28

P. A. ZIEGLER & F. ROURE: ALPINE OROGEN

NW

Belledonne

Po Plain SE

EUROPEAN

!S_cm/sr

^opemT,

40km

Apulian

Moho

Duplexes

made up

of European

lower crust

( Triangle zone)

Jura Mtns

Gran

Paradiso

Monferrato

Penninic

front

Ivrea

FIG. 4. Crustal section across the Western Alps.

Molasse

basin

l

0

10

20

30

40 -

50 -

60 km-

Heivetic

nappes

Aar

Massif

Penninic

nappes

Insubric

Line

©

Po Plain

FIG. 5. Crustal section across the Central Alps

Source : MNHN , Pahs

PERI-TETHYS MEMOIR 2: ALPINE BASINS AND FORELANDS

29

sequent imbrication of the European foreland crust,

resulting in the uplift of the external Pelvoux,

Belledonne and Mont Blanc crystalline massifs

(Fig. 4).

Nappes derived from the Apulian margin

occupy the northeastern part of the Western Alps

(Sesia-Lanzo and Dent Blanche nappes). The inter¬

nal Schistes Lustres nappe, containing ophiolites,

was derived from the Piemont Ocean. The Valais

Trough and the Brian^onnais were deformed into a

system of west- and east-verging nappes, respec¬

tively, which overrode the passive foreland margin,

causing late Eocene-Oligocene subsidence of a

flexural foreland basin. Subsequent intense cou¬

pling of the orogenic wedge with the foreland is

evident by compressional reactivation of Mesozoic

tensional faults on the distal foreland margin, the

step-wise imbrication of the foreland crust and

detachment of the passive margin series from their

basement. By this process the earlier developed

foreland basins was largely destroyed. During late

Eocene to Miocene times, the West-Alpine fore¬

land was transected by a system of grabens which

form part of the European Cenozoic rift system

(Ziegler, 1994b). During Miocene and Pliocene

times, Mesozoic and Cenozoic extensional basins

in the foreland of the Western Alps were inverted,

causing disruption of the Mesozoic proximal pas¬

sive margin sedimentary prism; moreover thrusting

propagate into the domain of the Jura Mtns. (Roure

et al., 1990; Roure and Colletta, Philippe et al, this

volume).

Commensurate with the dominantly west-ver-

gence of the external Western Alps, their up to

60 km deep crustal root (Laubscher, 1992b) is

located under their eastern, internal parts. More¬

over, seismic tomography images an about 175 km

deep lithospheric root, located beneath the western

Po Plain and an apparently detached, east-dipping

subduclion slab (Spakman, 1990). The Western

Alps are devoid of subduction related magmatism.

Their evolution reflects increasing collisional cou¬

pling of the orogenic wedge with the European

foreland in which contemporaneous inversion

structures developed as far away as in the Western

Approaches and the Celtic Sea (Ziegler, 1990;

Ziegler et al., 1995).

The external parts of the Western Alps and

their foreland contain no significant hydrocarbon

accumulations. Sub-thrust plays are areally restrict¬

ed due to the limited extent of the foreland beneath

the frontal thrusts.

Central Alps

The Central Alps evolved by Late Cretaceous

and Paleogene closure of the larger South Penninic

and the smaller North Penninic Valais oceanic

basins. The total amount of crustal shortening

achieved across the Central Alps is in the range of

500 to 550 km (Fig. 5; Schmid et al., 1996; Ziegler

et al., this volume).

The Central Alps are characterized by major,

north-verging, partly basement cored nappes. The

Austroalpine nappes were derived from the south¬

eastern margin of the South Penninic Ocean. The

Penninic nappes derive from the North and South

Penninic troughs and the Briantjonnais block, sepa¬

rating them. The sedimentary Helvetic nappes

were derived from the northern, European shelf;

their basement core is represented by the Gotthard

Massif and the lowermost Penninic nappes. Mio-

Pliocene uplift of the external Aar Massif, entailing

deformation of the overlaying stack of nappes, was

partly contemporaneous with the development of

the Jura fold-and-thrust belt. The Southern Alps

consist of a south-verging internal stack of base¬

ment imbrications and an external, thin-skinned

thrust belt.

End Cretaceous to early Paleocene closure of

the South Penninic trough was accompanied by

large radius deformation of the Helvetic shelf,

causing erosion of much of its Cretaceous sedi¬

mentary cover. Late Eocene closure of the North

Penninic trough was followed by flexural subsi¬

dence of the Helvetic shelf under the load of the

advancing Austroalpine, Penninic and Helvetic

nappes; the Helvetic nappes consist of sediments

which were detached from the foreland crust. Post-

collisional indentation of Apulia, amounting to

some 120 km, caused overthickening of the oro¬

genic wedge. This gave rise to Oligocene back-

folding and -thrusting along the Insubric line,

Mio-Pliocene step-wise imbrication of the northern

and southern foreland crust and ultimately thrust

propagation into the conjugate flexural foreland

basins, causing their partial destruction. Post-colli-

30

P. A ZIEGLER & F. ROURE: ALPINE OROGEN

sional crustal shortening was accompanied by the

subduction of continental lithospheric material.

South-directed underthrusting of the European

foreland gave rise to the development of a some

60 km deep crustal root beneath the southern part

of the Central Alps and a 175 km deep lithospheric

root located beneath the northern parts of the Fo

Plain (Spakman, 1990). Oligocene detachment of

an earlier formed subduction slab from the lithos¬

phere was accompanied by the intrusion of partial

melts derived from the mantle-lithosphere (von

Blankenburg and Davies, 1995).

The South-Alpine external, thin skinned thrust

belt evolved out of a Triassic extensional basin, a

superimposed passive margin prism and a late

Oligocene-mid-Miocene flexural foreland basin; it

hosts significant hydrocarbon accumulations

(Anclli et al., this volume). The northern, Molasse

foreland basin contains a thick syn-orogenic clastic

wedge which rests on a relatively thin passive mar¬

gin sequence. This basin is internally little

deformed, extends only some 25 km beneath the

Alpine nappes to the frontal basement imbrications

of the Aar massif and is limited to the north by the

Jura fold-and-thrust belt. The Molasse Basin has a

very limited hydrocarbon potential (Ziegler et al.,

Philippe et al., this volume).

Eastern Alps

In contrast to the Central Alps, the Eastern

Alps are characterized by an autochthonous base¬

ment which extends from the northern thrust front

for at least 60 and perhaps as much as 100 km

under the stack of Austroalpine nappes (Zimmer

and Wessely, Tari, this volume). The eastward dis¬

appearance of major Helvetic nappes, involving

European passive margin sediments, is striking. As

Penninic nappes play a subordinate role in the

architecture of the Eastern Alps, it is assumed that

the Valais Trough terminates to the east or merges

with the South Penninic Trough. In this context,

Froitzheim et al. (1996) propose that the crystalline

core of the Tauern window is formed by European

crust and, as such, is not equivalent to the Bri-

antjonnais, as commonly suggested (Tollmann,

1985).

Evolution of the Eastern Alps involved Late

Jurassic closure of the Hallstatt and Cretaceous

closure of the Penninic oceans. Cretaceous mass

transport was northwesterly directed. Profound late

Senonian and Paleocene disruption of the Euro¬

pean passive margin shelf, involving transpression-

al reactivation of Permo-Carboniferous and

NW

Outer

Pieniny

Klippen belt

Tatras

Inner West

Mtns Carpathians

Pannonian

basins

SE

FIG. 6. Crustal sections across the Carpathians, imaging the changes from colli¬

sion to back-arc extension or collapse (after Roure et al., 1996).

Source :

PERI-TETHYS MEMOIR 2: ALPINE BASINS AND FORELANDS

31

Mesozoic crustal discontinuities and ensuing uplift

and erosion (Ziegler, 1990), presumably reflects

initial intense coupling between the evolving oro-

genic wedge and its foreland. During the Paleo¬

gene the Austroalpine nappes were thrusted over

the foreland crust; this was accompanied by the

flexural subsidence of the eastern Molasse basin,

which is characterized by an array of syn-flexural

tensional faults. By early Miocene times, the sedi¬

mentary Austroalpine nappes had arrive near their

present location. During the emplacement of the

East-Alpine stack of nappes, the orogen was appar¬

ently mechanically decoupled from the foreland, as

indicated by the absence of compressional foreland

structures. Paleocene destruction of the Mesozoic

passive margin prism, particularly in areas located

to the south of the Bohemian massif, is held

responsible for the eastward disappearance of the

Helvetic nappes.

The small remnant foreland basin of the East¬

ern Alps hosts several petroleum systems. These

are tied to Mesozoic and early Oligocene oil

source-rocks and to biogenic gas generated in the

Oligocene and early Miocene deeper water clastic

series of the foreland basin fill. The subthrust play

of the Eastern Alps has to contend with consider¬

able reservoir and structural risks, deep objectives

and topographic difficulties. The sedimentary

allochthon contains important hydrocarbon accu¬

mulations beneath the Vienna Basin; elsewhere its

hydrocarbon potential has as yet to be proven

(Roeder and Bachmann, Zimmer and Wessely, this

volume).

Northern and Eastern Flysch Carpathians

The Northern and Eastern Carpathians consist

of a stack of sedimentary nappes, involving Early

Cretaceous to Miocene flysch series which are

thrusted over the foreland. The latter extends at

least some 75 km under this orogenic wedge,

which accounts for at least 250 km of shortening

(Fig. 6). The ophiolite bearing Pienniny Klippen

zone marks the internal boundary of the Flysch

Carpathians, in which the involvement of passive

margin sediments remains conjectural (Sandulescu,

1984; Roure et al., 1993; Bessereau et al., Sovchik

and Vul, Dicea, this volume).

The Flysch Carpathians evolved in response to

Mid-Cretaceous and Cenozoic south and westward

subduction of the oceanic Magura-Piennide basin

and eastward displacement of the intra-Carpathian

North Pannonian, Tisza and Dacides blocks; these

record Late Jurassic and Early Cretaceous orogenic

events related to the closure of the Vardar-Hallstatt

Ocean (Sandulescu, 1984; Csontos et al., 1992;

Stefanescu and Bakes, this volume).

The passive margin sedimentary prism of the

North- and East-Carpathians forelands was disrupt¬

ed by the late Senonian and Paleocene deep inver¬

sion of the Polish Trough (Ziegler, 1990) and the

Early Cretaceous inversion of the Dobrogea

Trough (Belov et al., 1987). These features appear

to link up beneath the Eastern Carpathians. Corre¬

spondingly, large parts of the autochthonous fore¬

land beneath the Flysch Carpathians lack a thick

passive margin prism which otherwise could have

been detached from its basement during the Late

Cretaceous and Cenozoic phases of the Carpathian

orogeny. Reflection-seismic data show that flexural

subsidence of the autochthonous foreland was

accompanied by major normal faulting. Moreover,

they image mild compressional deformation of the

autochthonous basement beneath the internal parts

of the Flysch Carpathian accretionary wedge; how¬

ever, major basement structures of the Aar Massif

type are lacking (Roure et al., 1993, Sovchik and

Vul, Dicea, this volume).

Roll-back of the subducted oceanic slab and

its dehydration is tracked by a chain of mid-

Miocene to Pliocene calc-alkaline volcanics, paral¬

leling the Klippen zone (Szabo et al., 1992). Deep

reflection profiles through the North-Carpathians

image the steeply south dipping European foreland

crust (Tomek, 1993). The absence of post-Pale-

ocene compressional foreland structures and the

presence of only minor compressional basement

structures beneath the Flysch Carpathians suggest

that coupling between the orogenic wedge and its

foreland was at a low level. Post-Oligocene devel¬

opment of the Flysch Carpathians was accompa¬

nied by the collapse of the Pannonian Basin in

their hinterland (Royden and Horvath, 1988; Tari

et al., 1992).

The Flysch Carpathians and their foreland

host important petroleum systems. The most

32

P. A. ZIEGLER & F. ROURE: ALPINE OROGEN

important one is related to the Oligocene Menilite

shales; a second potential petroleum system is

related to Cretaceous black-shales. Jurassic and

even Palaeozoic source-rocks have locally con¬

tributed hydrocarbons (Bessereau et al., Brzobo-

haty et al., Stefanescu and Baltes. Dicea, this

volume).

Peri-Adriatic Thrust Belts

The stable Adriatic (Apulian) platform is

flanked to the east by the Dinarides-Albanides and

to the west by the Apennines. Whereas in the

Dinarides-Albanides orogenic activity commenced

during the Late Jurassic and persisted variably into

Miocene and Pliocene times, the Apennines

evolved essentially during Neogene times (Frasheri

et al., Anelli et al., this volume).

The west-verging Albanides are characterized

by thin skinned thrust sheets which are detached

from their autochthonous basement at the level of

Triassic evaporites. These thrust sheets involve a

thick Triassic-Jurassic passive margin sequence,

dominated by carbonates, and Cretaceous to Paleo¬

gene flysch series which grade to the west into

platform carbonates. Mio-Pliocene molasse series

are involved in the most external zone. Westward

progression of the flysch facies through time

describes the gradual advance of the orogenic front

and the migration of the flexural foreland basin.

The up to 14 km thick Mirdita ophiolite nappe

forms the orogenic lid of the Albanides. These

ophiolites, which presumably represent the oceanic

crust of Sub-Pelagonian trough, were obducted

during the Late Jurassic closure of the Vardar sys¬

tem of oceanic basins. The autochthonous foreland

crust extends essentially unbroken from the thrust

front of the Albanides at least 100 km under their

internal nappes

The external, Ionian zone of the Albanides

host a major hydrocarbon province which derives

its charge from multiple Mesozoic source-rocks

(Frasheri et al., this volume).

In contrast to the long lived Albanides, the

Apennines are a young orogenic belt which

evolved only during late Oligocene to Plio-Pleis-

tocene times. The Apennines are characterized by

major sedimentary nappes; these involve Triassic

to Paleogene syn- and post-rift shallow-water and

pelagic series and Oligocene to Miocene flysch.

deposited in an eastward migrating foreland basin.

The flysch nappes are detached from their substra¬

tum. An ophiolitic nappe, or rather an ophiolitic

melange, is only locally preserved along the

Tyrrhenian coast (Liguride units). The external

flysch nappes override a thick parautochthonous

and autochthonous Triassic to Paleogene sedimen¬

tary sequence, consisting predominantly of plat¬

form and pelagic carbonates (Fig. 7).

Following Paleogene closure of the Alboran-

Ligurian-Piemont Ocean, the Apennines evolved in

response to Oligocene and younger westward

underthrusting and subduction of the Adriatic fore¬

land. The flysch facies mirrors the progressive

eastward migration of the flexural foreland basins

and the gradual incorporation of its proximal parts

into the evolving orogenic wedge (Ricci Lucchi,

1986). The most internal units of the Apennines

are formed by Ligurides, representing the sedimen¬

tary fill of the Ligurian-Alboran Basin, and the

basement involving Calabria-Peloritani nappes

which are of uncertain origin. The main elements

of the allochthon are derived from the distal parts

of the Apulian platform which were separated from

its main parts by the deep-water Lagonegro trough.

Emplacement of these sedimentary nappes was

accompanied by partial detachment of the

autochthonous sedimentary cover of the Apulian

platform, its imbrication and by in-sequence thrust

propagation into the Adriatic Sea (Anelli et al., this

volume). There is only little evidence for involve¬

ment of the Apulian crust in the orogenic edifice of

the Apennines (Ponziani et al., 1995); correspond¬

ingly, the orogenic wedge and the foreland crust

were largely mechanically decoupled during the

Apennine orogeny.

Substantial supra-crustal shortening, evident

in the Apennines, was apparently compensated by

the subduction of continental lithosphere, giving

rise to extensive magmatic activity along the west¬

ern margin of Italy from mid-Miocene times

onward. Steepening and roll-back of the subducted

lithospheric slab is held responsible for contempo¬

raneous back-arc rifting and the gradual opening of

the Tyrrhenian Sea (Spakman et al., 1993; Serri et

al., 1993; Doglioni, 1993).

Source : MNHN. Paris

PERI-TETHYS MEMOIR 2: ALPINE BASINS AND FORELANDS

33

Source : MNHN , Pahs

34

P A. ZIEGLER & F. ROURE: ALPINE OROGEN

The Apennines and their foreland basin host

important petroleum systems which are related to

Triassic and Jurassic source-rocks, deposited dur¬

ing the syn-rift stage. Secondary petroleum sys¬

tems are related to the passive margin sequence

and the syn-orogenic flexural basin. The latter con¬

tains major amounts of biogenic gas (Anelli et al.,

this volume).

East-Taurus Foothills Thrust Belt

The foothills of the eastern Taurus orogenic

belt consist of a 30 to 40 km wide zone character¬

ized by south-verging thin-skinned thrust imbri¬

cates, cored by Albian-Campanian shelf carbonates

which were detached from the autochthonous

Palaeozoic sedimentary cover of the Arabian

Shield. To the north this thrust belt is limited by

ophiolite and matamorphic nappes. Its southern

foreland is characterized by a broad belt of intra-

plate congressional structures. This thrust belt

developed during the Late Senonian to Eocene col¬

lision and of the North-Arabian passive margin

with the Taurides arc-trench system (Yilmaz, 1993;

Gilmour and Makcl, this volume).

The North-Arabian passive margin developed

during the Late Triassic with the opening of the

Tethys Ocean. Its passive margin sequence spans

Jurassic to Campanian times; it is interrupted by a

major Early Cretaceous hiatus which truncates

Jurassic strata to the north. Development of a flex¬

ural foreland basin commenced during the late

Senonian and was accompanied by the obduction

of the ophiolitic nappes and imbrications of the

Cretaceous passive margin series. The external

thin-skinned thrusted structures are sealed by Pale-

ocene and younger series which were gently folded

during the late Eocene. Eocene and Miocene con¬

tinued thrusting was confined to the internal zones

of the orogen and was accompanied by the closure

of remnant oceanic basins. Miocene(?) foreland

compression caused reactivation pre-existing

crustal discontinuities, resulting in the upthrusting

of foreland structures at distances of over 100 km

to the south of the Alpine trust front.

The external, thrusted structures of the Taurus

foothills are charged with hydrocarbons generated

by Silurian shales of the autochthonous Arabian

foreland, forming part of the Tethyan pre-rift

sequence (Gilmour and Make), this volume).

To round-off this pallet of diversified structur¬

al styles of the Alpine chains, attention is drawn to

the paper by J. Flinch (this volume) who describes

the gravitational collapse of the external accre¬

tionary prism of the Rif-Betic Cordillera arc which

spilled over the Atlantic continental margin of

Morocco and Iberia, forming the Gibraltar

allochthon.

ALPINE FORELANDS

During the Late Cretaceous and Cenozoic

phases of the Alpine orogeny, both the European

and the Africa- Arabian forelands record a

sequence of intra-plate deformations which caused

drastic changes in their structural and palaeogeo-

graphic evolution.

Micro-tectonic analyses indicate that the tra¬

jectories of principal horizontal compressional

stress axes affecting the European and Africa-Ara-

bian platforms changed repeatedly during Late

Cretaceous and Cenozoic times (Letouzey and Tre-

molieres, 1980; Bergerat, 1987; Muller, 1987; Bles

et al., 1989; Lacombe et al., 1990); this can be

related to changes in their convergence pattern.

Moreover, from late Eocene onward, gradual

development of the Red Sea-Suez-Libyan-Pelagian

Shelf and Rhine-Rhone-Valencia-Trans-Atlas rift

systems was broadly contemporaneous with com¬

pressional intra-plate deformations. This is inter¬

preted as reflecting the interference between stress

systems related to the late phases of the Alpine col¬

lision and stresses related the gradual assertion of a

new tensional kinematic regime, governing a fun¬

damental plate-boundary reorganization, which

may ultimately lead to the break-up of the present

continent assembly (Ziegler, 1988, 1990, 1994b;

Ziegler et al., 1995).

Source : MNHN, Paris

PERI-TETHYS MEMOIR 2: ALPINE BASINS AND FORELANDS

35

Intra-plate Compression

During the Late Cretaceous and Cenozoic,

intra-plate compressional stresses reactivated a

broad spectrum of tensional and transtensional

Mesozoic and Palaeozoic basins in Europe and in

Africa- Arabia, causing their inversion (Fig. 1). At

the same time there are indications for broad buck¬

ling of the lithosphere, such as Plio-Pleistocene

accelerated subsidence of the North Sea basin,

uplift of the Fennoscandian Shield and Neogene

isolation of the Paris Basin (Kooi et al., 1989;

Ziegler, 1990; Ziegler et al., 1995).

In Europe, inversion structures and upthrusted

basement blocks occur within a radius of up to

1500 km from the thrust front of the Alpine oro-

gen. Late Cretaceous and Paleogene inversion fea¬

tures are located in the northern and northwestern

foreland of the Alps and in the foreland of the

Pyrenees, whereas Eocene and younger inversion

structures are located in the foreland of the West¬

ern Alps. (Ziegler, 1987, 1990; Roure and Colletta,

this volume).

In northern Africa, main intra-plate compres¬

sional structures evolved by latest Cretaceous and

Eocene inversion of the Triassic Atlas rift system.

(Casero and Roure, 1994; Zizi, this volume). The

Saoura-Ougarta chain is a Palaeozoic rift which

was inverted during the Permo-Carboniferous and

underwent a second Alpine inversion during the

Eocene (Ziegler, 1988).

On the Arabian craton, the Sinai arc, consist¬

ing of the Negev and the Palmyrides fold belts,

represents a major inversion feature. The

Palmyrides developed by inversion of a Permo-Tri-

assic aborted rift. First mild compressional defor¬

mations occurred at the end of the Cretaceous; the

major Mio-Pliocene phase of basin inversion

involved transpressional deformations in response

to NNW-SSE directed compression. Inversion of

the Palmyra trough was contemporaneous with the

thrusting phases in the East-Taurus orogenic belt

and early movements along the Dead Sea wrench

fault (McBriden et al., 1990; Chaimov et al.,

1993). Transpressional deformation of the Negev

fold belt is mainly late Senonian in age; minor

inversion movements continued into the Miocene

(Quennell, 1984; Moustafa and Khalil, 1995).

Basically we recognize two phases of prevail¬

ing intra-plate compression. The first phase is

related to initial collisional coupling between a

nascent orogen and its foreland (e.g. Carpathian

and East-Alpine foreland). Such deformations may

be controlled by the crustal configuration of the

respective passive margin (upper or lower plate,

availability of a thick passive margin prism or the

lack thereof), the rheological characteristics of the

oceanic lithosphere separating the accretionary

wedge from the respective passive margin and the

rate of their convergence. The second phase of

foreland compression is related to lithospheric

overthickening of the orogenic wedge and to

resulting thrust propagation into the foreland (e.g.

West and Central Alpine foreland). Processes gov¬

erning the coupling and decoupling of an orogenic

wedge and its foreland are, however, still poorly

understood (Ziegler et al., 1995).

Inversion can have severe repercussions on

the hydrocarbon habitat of a basin, mainly by

reversing its subsidence pattern, by re-configura-

tion of pre-existing structural traps and profound

erosion, resulting in the loss of hydrocarbons to the

surface. However, partly inverted basins can host

important hydrocarbon provinces such as those of

the Lower Saxony and West Netherlands basins

(Ziegler, 1990, 1995b).

Cenozoic Rift Systems

Whereas lateral escape of, for instance, the

Intra-Carpathian blocks, rotation of microplates,

such as the Corsica-Sardinia block, and progres¬

sive retreat of subducted slabs in the Tyrrhenian

Sea and the Aegean arc controlled the distribution

of Neogene extensional structures within the frame

of the overall compressional Alpine system of

fold-and-thrust belts, late Eocene and younger

development of the Rhine-Rhone and the East-

African-Gulf of Suez-Libyan-Pelagian Shelf rift

systems cannot be directly linked to the evolution

of the Alpine orogen (Fig. 1).

The Cenozoic rift system of Western and Cen¬

tral Europe extends from the shores of the North

Sea over a distance of some 1 100 km into the

West-Mediterranean domain; from there an alka-

36

P. A. ZIEGLER & F. ROURE: ALPINE OROGEN

line volcanic chain projects southwestwards across

the Alboran Sea, the Rif fold belt and the Atlas

ranges to the Atlantic coast and to the Cape Verdes

Islands. Including this volcanic chain, the entire

rift system has a length of 3000 km. This rift sys¬

tem began to evolve during the middle and late

Eocene in the European Alpine foreland and prop¬

agated during the Oligocene northward and south¬

ward. In the western Mediterranean, crustal

extension culminated, after a rifting period of only

7 Ma, in Miocene crustal separation and the open¬

ing of the oceanic Algero-Provengal Basin; this

involved a counter-clockwise rotation of the Corsi-

ca-Sardinia block (Torne et al., Vially and Tre-

molieres, this volume). During Miocene and

Plio-Pleistocene times, tectonic and partly also vol¬

canic activity persisted along the different seg¬

ments of this mega-rift system, albeit under

changing regional stress regimes. Development of

the Cenozoic rift system of Western and Central

Europe was contemporaneous with the Eocene and

later phases of the Alpine orogeny, during which

the northwestern Alpine foreland was repeatedly

subjected to horizontal intra-plate compressional

stresses, causing inversion of Mesozoic extensional

basins at considerable distances from the Alpine

thrust front. The southern elements of the Euro¬

pean Cenozoic rift system cross-cuts the Alpine

chains of the western Mediterranean domain.

Viewed on a broader scale, evolution of the

West and Central European Cenozoic rift system

was broadly contemporaneous with the develop¬

ment of the East African-Red Sea, Libyan and

Pelagian Shelf rift systems; its Neogene develop¬

ment was paralleled by back-arc extension govern¬

ing the subsidence of the Pannonian Basin and the

Aegean, Tyrrhenian and Alboran seas. As such the

Rhine-Rhone-Valencia and the Red Sea-Libyan-

Pelagian Shelf rift systems can be considered as

forming part of the Neogene Alpine-Mediterranean

collapse system (Ziegler, 1988, 1990). In this con¬

text, Neogene progressive eastward escape of Apu¬

lia and Sicily relative to North African may have

contributed to the evolution of the Pelagian rift

system (Casero and Roure, 1994).

However, considering the dimensions of the

West European-West African and the East-African-

Red Sea-Libyan rift systems, it is hardly conceiv¬

able that such major rifts developed solely in

response to the Alpine collision. More likely, they

form part of a new break-up system which may

culminate in the disruption of the Alpine plate

assembly (Ziegler, 1994b).

The European Cenozoic rift system hosts the

Rhine Graben and Valencia Trough hydrocarbon

provinces; whereas the former relies exclusively on

syn-rift source-rocks, the latter relies on a combi¬

nation of syn- and pre-rift source-rocks (Ziegler,

1995b, Torne et al, this volume). The major Gulf of

Suez hydrocarbon province relies exclusively on

Late Cretaceous pre-rift source-rocks (Ziegler,

1995b). The hydrocarbon charge of the Pannonian

system of back-arc basins is related to a combina¬

tion of syn- and pre-rift source-rocks. Uplift of

major rift domes, such as those associated with the

Rhine Graben and the Red Sea, causes disruption

of the pre-rift platform sedimentary sequences and

changes the hydrodynamic setting of the remnant

rift flank basins (e.g. Paris Basin).

PETROLEUM SYSTEMS OF ALPINE

FORELANDS AND EXTERNAL THRUST

BELTS

The complex geodynamic and tectonic evolu¬

tion of the Peri-Tethyan platforms and the

subsequent Alpine development of the European-

Africa-Arabian plate boundaries have conditioned

both the distribution and maturation history of

potential source rocks in the conjugate forelands as

well as within the Alpine orogen and its successor

basins.

Pre-Orogenic Versus Syn-Orogenic Source

Rocks

Palaeozoic source-rocks of the pre-rift

sequence, preserving part of their initial petroleum

potential until the onset of Alpine flexuring and

thrusting, play an important role on the Arabian

platform (Silurian shales; Gilmour and Miikel, this

volume) and in Western and Central Europe (main-

Source : MNHN. Paris

PERI-TETHYS MEMOIR 2: ALPINE BASINS AND FORELANDS

37

ly Carboniferous and Autunian; Ziegler, 1990).

Numerous Mesozoic source rocks have been iden¬

tified within the Mesozoic syn-rift series (mainly

Triassic and Early Jurassic) and in the passive mar¬

gin sequences of the former Tethys (mainly Late

Jurassic to late Early Cretaceous). These pre-oro-

genic sequences account for most of the oil and

thermal gas potential of the Alpine orogen and its

associated basins.

In addition, organic-rich series (TOC values

up to 10%) have been identified in the Oligocene

syn-flexural fill of the East-Alpine and Carpathian

foredeep (Menilite series); these are coeval with

the prolific Maykop source-rocks of the Black Sea,

Crimea and the foredeeps of the Great Caucasus

and Southern Caspian. Based on oil/source-rock

correlations and the occurrence of the oleanane

biomarker, even in subthrust Mesozoic reservoirs

of the foreland, these Oligocene series are believed

to be the most efficient source-rock in the Outer

Carpathians and the adjacent East-European fore¬

land, from Poland to Romania (Ulmishek and

Klemme, 1990; Koltun, 1992; Ten Haven et al.,

1993; Lafargue et al., 1994; Bessereau et al., this

volume). Similar facies occur in the German and

Austrian Molasse basin where they are the primary

source of accumulated oils (Zimmer and Wessely,

Roeder and Bachmann, this volume). Oligocene

series display also good TOC values in the succes¬

sor Pannonian basin where they have been suffi¬

ciently buried, under a relatively high geothermal

gradient, to enter the oil window and account for

significant hydrocarbon reserves in Hungary and

former Yugoslavia. In addition, Oligocene shales

have a good source-potential in the Cenozoic rift

basins of Western Europe (Vially and Tremolieres,

Torne et al., this volume)

The thick Mio-Pliocene terrigenous fill of the

Carpathian and Periadriatic foredeeps (Apennines,

Albanides) have TOC values averaging 1%. Most¬

ly immature, these Neogene series account for

large reserves of bacterially-generated gas, but

only for minor oil (Kotarba, 1992; Anclli et al., this

volume).

Sedimentary Versus Tectonic Burial of Source-

Rocks

For most Alpine foreland fold-and-thrust belt

systems, available surface and sub-surface geologi¬

cal and geophysical data provide sufficient con¬

straints to construct reliable regional structural

cross-sections, extending from the autochthonous

foreland to the frontal parts of the thrust belt. Com¬

bined with a good biostratigraphic dating of the

syn-orogenic sequences, this permits construction

of high quality balanced cross-sections and their

step-wise palinspastic restoration, retracing the

evolution of the respective area from the onset of

deformation to its Present configuration (Bally et

al., 1988; Roure et al., 1993; Zoetemeijer et al.,

1993).

Forward kinematic modelling permits to

retrace the maturation history of potential source-

rocks within an evolving Alpine foredeep basins,

i.e. during their initial sedimentary burial and their

subsequent tectonic burial beneath the advancing

allochthonous units. Ultimately, also uplift and ero¬

sion, related to tectonic accretion of individual

units into the Alpine orogenic wedge, can be simu¬

lated. Coupled with palaeo-thermal reconstructions

and transformation kinetics of kerogens into hydro¬

carbons, these new techniques provide an efficient

tool to date the timing of source-rock maturation

and to evaluate migration pathways between

hydrocarbon kitchens and potential traps (Roure

and Sassi, 1995).

Distinct source-rocks frequently coexist in a

single Alpine foreland basin. However, due to lat¬

eral and vertical changes in their distribution, they

usually contribute to independent petroleum sys¬

tems, characterized by unique timing of matura¬

tion. hydrocarbon expulsion and migration.

Therefore, it is essential to obtain an impression of

potential migration pathways between source areas

and traps and the timing of trap development,

before drilling a prospect, in order to decrease

exploration risks in such complex areas as the

Alpine orogen.

38

P. A. ZIEGLER & F. ROURE: ALPINE OROGEN

FIG. 8. Evolution of a foreland fold-and-thrust belt system and distribution of

effective petroleum systems:

a) Sedimentary burial, early maturation and long-range migration pathways

b) Tectonic burial, late maturation and short range migration pathways.

Long-Distance Lateral Versus Short-Distance

Vertical Migration

Calibration and reconstruction of palaeo-tem-

peratures and -hydrodynamics (convective heat

and fluid transfers) has hardly been addressed for

the external imbricate structures of the different

Alpine fold-and-thrust belts. However, distinct

episodes of petroleum generation and migration are

usually recognized in the Carpathians, Apennines

and Albanides, which constitute the most produc¬

tive petroleum provinces in the European part of

the Alpine orogen (Roure and Howell, 1996).

Most frequently, oil is generated early in fore¬

land basins, resulting in long range migration of

the first hydrocarbon products from the foredeep

source area toward the foreland, up-slope the

regional flexure. This buoyancy driven mechanism

is even facilitated by regional hydrodynamics,

accounting for recharge of meteoric waters in the

foothills and discharge in the foreland in the area

of a potential forebulge, if present (Fig. 8). Thus,

early long-distance updip migration of oils generat¬

ed by Triassic source-rocks in the Albanian fore¬

deep, charging the Aquila field, located on the

eastern Puglia slope in the Adriatic (Anelli et al.,

this volume), is indicated. Apparently, the Meso¬

zoic reservoirs at Lopushnia oil field in the Ukrain¬

ian sub-thrust foreland (Izotova and Popadyuk;

Sovchik and Vul, this volume), were laterally

charged by the Oligocene series of the Outer

Carpathians; these are presently entirely detached

Source :

PERI-TETHYS MEMOIR 2: ALPINE BASINS AND FORELANDS

39

from their former substratum and are now accreted

to the allochthon (Lafargue et al., 1994).

A residual oil potential is eventually preserved

within the allochthonous units or in the under-

thrusted foreland, accounting for late phases of

maturation and migration. Under such conditions,

expelled hydrocarbon can hardly reach the foreland

as frontal antiform structures are already well

developed and provide traps; instead, hydrocarbons

generated during such late stages are preferentially

trapped in nearby structures, requiring short dis¬

tance lateral and/or vertical migration. This con¬

cept applies, for instance, for the deeply buried

duplexes of the Borislaw-Pokut zone of the

Ukrainian Outer Carpathians (Sovchik and Vul,

this volume) and the subthrust para-autochthonous

plays of the Southern Apennines (Casero et al.,

1991; d' Andrea et al., 1993; Anelli et al, this vol¬

ume).

On the other hand, hydrocarbons generated by

source-rocks forming part of the under-thrusted

sedimentary cover of the foreland may migrate

vertically into the allochthon or into neo-

autochthonous basins (Fig. 8). An example of such

a “plumbing system" are the oil and gas accumula¬

tions in the Neogene fill of the Vienna Basin and in

the underlying Austroalpine allochthon; these were

charged hydrocarbons generated by autochthonous

Late Jurassic basinal shales (Ladwein et al., 1991,

Zimmer and Wessely, this volume). Similarly, the

structures of the East-Taurus foothills thrust belt

were charged by Silurian source-rocks of the

autochthonous foreland sedimentary cover

(Gilmour and Makel, this volume).

CONCLUSIONS

Opening of the different constituent oceanic

basins of the Mesozoic Tethys was controlled by a

sequence of rifting cycles, the kinematics of which

changed with the progressive development of dis¬

crete plate boundaries first between Gondwana and

Laurasia and later between Eurasia and Laurentia-

Greenland. During these opening phases of Tethys

and the Atlantic, a sinistral translation between

Africa-Arabia and Europe and the interaction with

intervening microcontinents, such as the Italo-

Dinarides block, is held responsible for the devel¬

opment of First subduction zones and the onset of

the Alpine orogenic cycle. With the Senonian onset

of convergence of Africa-Arabia and Europe, new

subduction zones developed and progressive clo¬

sure of the different Tethys oceanic basins was fol¬

lowed by multiple and diachronous collisional

events. However, the preservation of remnant

Tethys oceanic basins in the Ionian Sea and the

Eastern Mediterranean shows, that continent-to-

continent collision has not yet occurred along the

entire Alpine-Mediterranean suture of Africa-Ara¬

bia and Europe. Moreover, Neogene continental

rifting, lateral block escape and subduction slab

roll back has governed the opening of the oceanic

Algero-Proven?al and the partly oceanic Tyrrhen¬

ian basins.

The different fold-and-thrust belts of the

Alpine orogenic system display highly variable

architectures. These are partly controlled by the

intensity of collisional coupling between the

respective orogenic wedge and its cratonic foreland

(and hinterland), and partly by the availability of a

thick passive margin sedimentary wedge, or its

absence. Imbricated passive margin wedges char¬

acterize, for instance, the external elements of the

Apennines, the Dinarides-Albanides and the East-

Taurides. In contrast, major nappes, derived from

the southern margin of the South Penninic Ocean,

directly override the undeformed foreland of the

Eastern Alps which extends some 100 km beneath

these nappes. In the Central and Western Alps,

increasing collisional coupling between the fore¬

land and the evolving orogenic wedge resulted in

imbrication of the foreland crust, uplift of external

crystalline massifs and deformation of the overlay¬

ing stack of nappes. In contrast, the Flysch

Carpathians consist of an accretionary wedge

which was thrusted over the essentially unde¬

formed autochthonous foreland.

Most of the Alpine fold-and-thrust belts are

associated with more or less pronounced flexural

foreland basins; however, in some cases these were

destroyed by late thrust propagation into the fore¬

land. The Alpine foreland basins display major

variations in their their architecture and the facies

development of their sedimentary fill. An array of

minor syn-flexural normal faults characterizes the

40

P. A. ZIEGLER & F. ROURE: ALPINE OROGEN

German and Austrian Molasse Basin (Roeder and

Bachmann, Zimmer and Wessely, this volume). In

contrast, the Carpathian foreland basins are charac¬

terized by a limited number of major syn-flexural

faults. On the other hand, thrust propagation into

the foreland plays an important role in the Apen-

nine foredeep (Anelli et al., this volume). Whereas

the Swiss Molasse Basin is characterized by shal¬

low marine and continental elastics, lacking effec¬

tive reservoir/seal pairs (Ziegler et al, this volume),

the Molasse Basin of Austria contains Oligocene

and early Miocene deeper water elastics and

shales, grading upwards into deltaic series, which

provide for several reservoir-seal pairs (Zimmer

and Wessely, this volume). Similarly, the flysch-

type sediments of the Apennine foreland basins,

which grade up into a Pleistocene deltaic complex¬

es, contain multiple effective reservoir/seal pairs.

In contrast, the sedimentary fill of the East-Taurus

foreland basin consists of marls, evaporites and

carbonates (Gilmour and Makel, this volume).

Compressional intra-plate deformations char¬

acterize both the European and the Africa-Arabia

platform. These developed in response to collision-

al coupling of these forelands with the Alpine oro-

gen. Early collisional foreland deformations occur

in the Carpathian and East-Alpine forelands, where

they caused partial destruction of the passive mar¬

gin sedimentary prism and inversion Mesozoic ten-

sional basins and the upthrusting of basement

blocks as far north as Denmark and southern Swe¬

den. Late collisional foreland deformations occur,

for instance, in the forelands of the Western and

Central Alps and the Pyrenees. Examples of thin-

skinned foreland fold-an-thrust belts are the Lom¬

bardian belt of the Southern Alps (Ziegler et al.,

this volume), the Jura Mountains of Switzerland

and France (Philippe et al, this volume) and the

Prerifain chains of Morocco (Zizi, this volume).

Their development caused partial destruction of

pre-existing syn-orogenic flexural basins. Base¬

ment involving up-thrusted blocks characterize the

Bohemian Massif in the foreland of the Eastern

Alps (Ziegler, 1990). Inversion of major Mesozoic

grabens in Europe, in Iberia and on the African-

Arabian Platform probably involved the entire

crust and possible also the mantle lithosphere

(Ziegler et al., 1995).

Eocene and later development of the Rhine-

Rhone-Valencia and East Africa-Red Sea-Libyan-

Pelagian Shelf rift systems caused disruption of the

sedimentary cover of the European and Africa-

Arabian platforms, particularly across major ther¬

mal domes. This had repercussions on the

hydrodynamic conditions in remnant platform

basins and their hydrocarbon habitat.

The petroleum systems of the external parts of

the Alpine orogen, its foreland and internal basins

are highly variable. They can rely either on pre-rift,

syn-rift or syn-flexural source-rocks or a combina¬

tion thereof. Some of these source-rocks attained

maturity already during the passive margin or flex¬

ural basin stage; others reached maturity only dur¬

ing the emplacement of allochthonous units. Each

basin has its own case history.

The petroleum exploration history of the

Alpine orogen and its associated basins com¬

menced in the early Nineteenth century. First

hydrocarbon exploration efforts were based on sur¬

face seeps, concentrated in the shallow, frontal

anticlines of the Outer Carpathians of the Ukraine

and Romania. In the course of time, improved geo¬

physical methods provided new tools for imaging

sub-surface structures. This resulted in new discov¬

eries of oil and biogenic gas in such foreland

basins as the Po Plain, the Adria, the German and

Austrian Molasse Basin and in the Aquitaine

Basin, as well as in the Pannonian and Vienna neo-

autochthonous basins. With the development of

static corrections and the development of CDP-

methods, and ultimately of 3D-seismic, imaging of

even structurally complex areas has become possi¬

ble, thus providing access to complex thrusted

structures (Le Vot et al., this volume) and subthrust

prospects of the Lopushnia (Carpathians) and

Tempa Rossa type (Southern Apennines; Roure

and Sassi, 1995; Anelli et al., this volume).

No doubt, future exploration efforts in the

external parts of the Alpine orogen an beneath the

Neogene fill of the Pannonian Basin will yield fur¬

ther oil and thermal gas discoveries, provided it

can be established that efficient petroleum systems

are available, have survived the most recent defor¬

mations and that only minor hydrocarbon re-migra¬

tion has occurred.

Acknowledgments - The authors of this paper

wish to extend their thanks to Dr. A. Mascle (IFP)

and to Prof. S.M. Schmid ( University Basel) for

their constructive comments on an earlier version

Source :

PERI-TETHYS MEMOIR 2: ALPINE BASINS AND FORELANDS

41

of this manuscript. The Institut Frangais du Pet-

role is thanked for taking care of draughting the

text figures.

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Tectonic evolution and paleogeography of Europe

P. O. Yilmaz, I. O. Norton, D. Leary & R. J. Chuchla

Exxon Production Research Co, PO Box 2189,

Houston, TX 77252-2189, USA

ABSTRACT

Multiple rifting and suturing events through

Phanerozoic times amalgamated Europe as we

know it today. Our detailed analysis of the crustal

blocks, now forming of Europe, during the Cale¬

donian, Hercynian and Alpine orogenies, allowed

us to understand the influence of these events on

the hydrocarbon systems of Europe.

To summarize this, we present a series of 1 1

palaeogeographic maps from Carboniferous to

Pliocene times. These maps were produced as part

of a project to develop basin-wide models for

regional play element distribution in the major

hydrocarbon-producing basins of Europe.

Description of the tectonic evolution of

Europe can be divided into four main phases which

are related to motions between Baltica, North

American/Greenland and Gondwana. The first

phase culminated in the assembly of Laurussia

(Europe and North America/Greenland) during the

Early Palaeozoic Caledonian Orogeny; it was fol¬

lowed by the Carboniferous assembly of Pangea

(Laurussia and Gondwana) during the Hercynian

orogeny. The third phase, involving rifting and

separation of these blocks, started in Permian time.

The fourth and final phase, that continues today, is

the Alpine orogenic cycle which resulted from

convergence of Africa and Europe.

INTRODUCTION

We present a series of palaeogeographic maps

which summarizes our understanding of the geo¬

logic evolution of Europe since Carboniferous

times. These maps were produced as part of an

Exxon project to develop basin-wide models for

regional play element distribution in the major

hydrocarbon-producing basins of Europe.

The tectonic evolution of Europe can be divid¬

ed into four main phases which are related to

motions between Baltica, North American/Green¬

land and Gondwana. The first phase involved the

formation of Laurussia (Europe and North Ameri¬

ca/Greenland) during the Early Palaeozoic Cale¬

donian Orogeny. The second phase was the

Carboniferous assembly of Pangea (Laurussia and

Gondwana) during the Hercynian orogeny. The

Yilmaz. P. O., Norton, I. O., Leary, D. & Chuchla. R. J.. 1996. Tectonic evolution and paleogeography of Europe. In. Ziegler.

P. A. & Horvath, F. (eds), Peri-Tethys Memoir 2: Structure and Prospects of Alpine Basins and Forelands. Mem. Mus. natn. Hist, nat.,

170: 47-60 + Enclosures 1-13. Paris ISBN: 2-85653-507-0.

This article includes IS enclosures.

Source :

48

P. O. YILMAZ ET AL.. TECTONIC EVOLUTION AND PALEOGEOGRAPHY

third phase was dominated by rifting and separa¬

tion of these blocks, starting in Permian times. The

fourth and final phase continues today and corre¬

sponds to the Alpine orogenic cycle which results

from convergence of Africa and Europe.

For times younger than Jurassic, relative

motions of cratonic blocks were determined from

sea floor spreading data in the Atlantic. Motions

between Europe and Africa were essentially strike

slip from 180 to 1 10 Ma, then swung round to the

convergent motion which continues today. Pre-

Jurassic relative plate motions were determined

from a combination of palaeomagnetic data and

geologic data on the timing of tectonic events. An

important final constraint was that the derived rela¬

tive motions were required to produce a pattern

that was geologically reasonable, i.e. no conver¬

gent and divergent rates that exceed rates known

from global post-Jurassic plate motion rates, and

relative motion directions that agreed with the tec¬

tonic data. This last constraint is particularly

important in the Hercynian orogeny, which

includes significant amounts of strike slip motions.

PALAEOZOIC CRUSTAL BLOCKS

Palaeozoic crustal blocks, as they were assem¬

bled in Permian times at the end the Hercynian

orogeny, are shown in Plate 1. Brief descriptions of

these blocks are given below.

Baltica

Baltica, also known as the Russian Platform,

is the Precambrian core of Europe; it consists of

several Archean age blocks that were amalgamated

into cratonic Baltica before 1.6 Ga (Zonenshain et

al., 1990). Baltica is bounded on the west by the

Iapetus suture and on the east by the Ural suture.

The northern edge of Baltica we take to be the

Timan Belt and its extension along the northern

coast of Scandinavia. This boundary was reactivat¬

ed during the Late Cambrian Fenno-Scandian

orogeny. The southern boundary is less well

defined. In Early Palaeozoic time, Baltica faced the

Tomquist Sea to the (present day) south. The Torn-

quist Line itself, however, does not mark the edge

of Baltica; the edge is further outboard, buried

beneath younger cover (Cocks and Fortey, 1982).

Pechora

This is the crust under the Barents Sea and

includes Svalbard. Genesis of this area is poorly

understood; we assume it to be amalgamated by

the end of the Caledonian orogeny, but most of this

block probably consists of older crust.

Laurentia

North America and Greenland are Precambri¬

an cratons that make up the Laurentian block. Like

Baltica, Laurentia consists of Archean terranes that

were amalgamated during Precambrian times, cul¬

minating in the Grenville orogeny between 800

and 1000 Ma.

Avalonia

Avalonia consists of southern England, Ireland

and the northeastern seaboard of North America. It

rifted away from Gondwana during the Early Cam¬

brian (550 Ma) and was sutured to Laurentia dur¬

ing the Caledonian orogeny (first collision at

425 Ma, end of orogeny at 405 Ma; McKerrow,

1988). There is not enough reliable palaeomagnetic

data from Avalonia to determine its positions

between rifting and collision. A motion path for it

was determined by first plotting the positions of

Europe, North America and Africa at the start and

end of its motion, then interpolating positions

between these times so that the motion of Avalonia

was continuous.

Armorica

Armorica is the name for the western Iberian

Peninsula and also for what is now western and

northern France. Armorica is separated from the

rest of France and Iberia (the Southern European

Block) by a suture zone of Hercynian age. Struc¬

tural studies within Armorica indicate that it was

severely deformed during the Hercynian orogeny

by its collision with the Southern European Block

(Matte, 1986), which acted as a solid indentor,

wrapping Armorica around itself. Maximum com¬

pression of Armorica was 500 km and the axis of

maximum deformation forms a line of weakness

along which the Bay of Biscay developed in Late

Cretaceous times.

Source :

PERI-TETH YS MEMOIR 2: ALPINE BASINS AND FORELANDS

49

Southern Europe

Southern Europe consists of northeastern

Iberia, the Balearics, southern France, Corsica and

Sardinia and probably some of the Palaeozoic

crustal elements involved in the Alps. Like Armor¬

ica and Avalonia, this block consists of Pan

African affinity crust which rifted off Gondwana

during the Early Palaeozoic.

Rhenohercynian

We interpret the Rhenohercynian zone as a

zone of Caledonian accretion that was the locus of

Devonian extension, as evidenced by the occur¬

rence of bimodal volcanics of Devonian age.

Bohemian Massif

The Bohemian Massif consists of a Precam-

brian terrane which, according to deep seismic

data, must be separated from the Bmo-Malopolska

Block (Suk et al., 1984).

Brno-Malopolska

The Brno-Malopolska Block is composed of

Precambrian granites and metamorphic rocks. This

unit includes the southern Holy Cross Mountains

area (Malopolska Massif).

Moesia

This block is assumed to consist of some Pre¬

cambrian crust that was accreted to southern Balti-

ca during the Early Palaeozoic.

Tisza

This block includes crust with European affin¬

ity (Royden and Baldi, 1988). It rifted from Europe

in Jurassic time, then joined Apulia in colliding

with Europe during the Alpine orogeny. Tisza's

crystalline and Mesozoic rocks outcrop only near

its eastern and western terminations.

PALAEOGEOGRAPHIC MAP FORMAT

The palaeogeographic maps presented here

were designed to show depositional environments

using the following colours:

Dark brown: highlands, considered to be sedi¬

ment source areas.

Light brown: lowlands, or zones of sediment

Pink and red colours distinguish collision- and

extension-related igneous rocks. The only sedi¬

mentary lithology shown is a chevron pattern for

evaporites. All active structures are indicated using

standard symbols as shown on the map legends.

Dashed lines show some political boundaries, pre¬

sent-day coastlines are in blue and some geograph¬

ic zones are identified with letter codes. For

orientation purposes, some cities are also shown.

Maps are plotted on an Albers equal area projec¬

tion (standard parallels 44° and 67°) with Europe

in its present-day position. A 5° present-day lati¬

tude/longitude grid is included, as are palaeolati-

tude lines derived from a compilation of

palaeomagnetic data.

PALAEOZOIC PALAEOGEOGRAPHY

Mid-Carboniferous (Namurian, 322 Ma)

The Mid-Carboniferous (Namurian, Ser-

pukhovian) map, Plate 2, illustrates collision of

Armorica and the South-European Block near the

end of the Hercynian orogeny. This collisional and

magmatic episode was largely ensialic, and result¬

ed in emplacement of abundant synorogenic gran¬

ites, shown in pink. Widespread orogenic

deformations occurred across northern Europe and

Iberia and in the Armorican, Saxothuringian,

Source

50

P. O. YILMAZ ET AL.: TECTONIC EVOLUTION AND PALEOGEOGRAPHY

Bohemian, Silesian, Massif Central, Ligerian, Cor-

sica-Sardinia and Carnic Alps areas.

Continental elastics were deposited in the

Armorican and Saxothuringian basins; linear

basins formed within the collision zone. Flysch

was deposited in foredeeps on either side of the

main orogenic belt. Principal flysch basins are the

Cantabrian Basin in the south, and the Rhenish

Basin in the north. The Rhenish Basin was eventu¬

ally Filled to capacity, and by Late Carboniferous

time (next time slice) marine connections to this

basin were severed.

Upper Carboniferous (Westphalian A/B,

306 Ma)

Hercynian deformation continued into the

Late Carboniferous. This final phase of crustal

shortening is sometimes referred to as the Variscan

phase. Plate 3 shows palaeogeography for the

Westphalian A/B stage of the Late Carboniferous.

Widespread Variscan deformations consist of

thrust- and wrench-faulting, folding, post-tectonic

granite emplacement and the accumulation of thick

continental sediments in the developing foredeeps.

Marine connections to the North-European fore¬

deep basin, located along the northern flank of the

Hercynian orogenic belt, were cut off as it progres¬

sively filled with elastics. In this basin, thick coal

measures were deposited during Westphalian

times. These provide the source for most of the gas

found in the Rotliegendes sandstones of the South¬

ern Permian Basin. Sedimentation on the South

Apulian shelf was locally disrupted by Variscan

tectonic events.

Lower Permian (Rotliegendes, 254 Ma)

Lower Permian time was dominated by the

collapse of the Hercynian mountain ranges and the

deposition of thick clastic sequences in the area of

their northern foreland basin. The Rotliegendes

(Plate 4) clastic reservoir facies was the first

sequence to be deposited. Sedimentation was

accommodated partly by thermal subsidence and

partly by continued subsidence of the relict

Variscan foredeep. This basin is known as the

Southern Permian Basin. Possible dextral shear

between Gondwana and Europe created intraconti¬

nental transform systems creating local transten-

sional and pull-apart basins. Individually, these

faults show relatively small displacements.

Grabens along the faults filled with continental

elastics. Marine shelf sedimentation continued in

the Apulian area.

Upper Permian (Zechstein, 251 Ma)

In Upper Permian time (Plate 5) a marine con¬

nection was established between the Arctic shelves

and the Northern and Southern Permian basins via

the Arctic-North Atlantic rift system (Ziegler,

1988). In the extensive Zechstein inland sea,

glacio-eustatic cycles controlled the accumulation

of alternating carbonate and evaporite deposits.

Late Permian fauna suggest communication

between the Boreal Zechstein seas and the Tethys

seas via Dobrudgea. Apart from providing haloki-

netically induced structural traps, the Zechstein

evaporites are an important seal facies, which seals

the Rotliegendes sands. Further to the south, on the

Apulian platform, rifting created basins containing

both evaporitic and continental deposits.

MESOZOIC CRUSTAL BLOCKS

Crustal blocks involved in the Mesozoic and

Cenozoic Alpine-Carpathian deformation are

shown on a Present-Day base map in Plate 6.

Europe and Africa remained relatively stable

through this time, although older structural grain

was reactivated during the Alpine orogeny. Europe

and Africa are relatively stable plates to which dif¬

ferent crustal blocks were amalgamated through

Mesozoic and Cenozoic times. Amalgamation is

Source :

PERI-TETHYS MEMOIR 2: ALPINE BASINS AND FORELANDS

51

still going on today with active subduction at the

Hellenic trench.

Iberia

The Iberian Block behaved independently of

Europe and Africa during the Late Cretaceous and

Paleogene (Choukroune et al., 1989; Roure et al.,

1989). The cratonic core of the Iberian peninsula

consists of Precambrian and Palaeozoic rocks

which have undergone Hercynian structuring dur¬

ing the Late Palaeozoic (Pin, 1990; Franke and

Engel, 1986; Ziegler, 1988, 1990).

Apulia

Apulia played a key role in the Alpine-

Carpathian deformations (Channell et al., 1979;

Biju-Duval et al., 1977). The Apulian Block

extends from Italy through the Pannonian area, the

Adriatic, Greece and further east to Turkey. There

is no known pre-Hercynian crust in Apulia. It

probably formed as accretionary crust during the

Hercynian collision of Gondwana with Europe.

Apulian crust was extensively affected by subse¬

quent Mesozoic rifting, until Jurassic separation

from Africa eventually led to formation of the

Eastern Mediterranean oceanic crust. In Tertiary

times, Apulia collided with Europe, initiating the

Alpine-Carpathian orogen. During this event, the

Apulian Block was shortened and partly subduct¬

ed. East-dipping subduction formed the Dinaric

belt along the eastern margin of Apulia, and west¬

dipping subduction formed the Apennine belt

along its western margin (Royden, 1993; Doglioni,

1992; Casero et al., 1990; Moretti and Royden,

1988; Frasheri et al., Anelli et al., this volume).

Tisza and Moesia

Tisza and Moesia were discussed above in the

section on Palaeozoic blocks.

Sakarya

Sakarya is a Cimmerian block which rifted off

the northern margin of Gondwana during Late Per¬

mian to Triassic times and collided with Europe in

Late Triassic to Liassic times (Sengor, 1984; Sen-

gor et al., 1984).

Rhodope

Rhodope is part of Europe. It rifted off Europe

but never strayed very far before colliding with the

European margin during the Meso-Alpine orogeny

(Dixon and Dimitriadis, 1984; Sengor, 1984).

MESOZOIC PALAEOGEOGRAPHY

Upper Triassic (Rhaetian, 210 Ma)

Plate 7 shows Upper Triassic palaeogeogra-

phy. Triassic times were characterized by regional

extension with multidirectional systems of grabens

being superimposed on Hercynian structural

trends. This extension, known as the Tethyan rift

event, had a profound influence on hydrocarbon

plays in the Apulian and Carpathian regions. Shal¬

low water reefal limestones characterize rifted

margins throughout Apulia. Rifting was accompa¬

nied by a widespread extrusive and shallow intru¬

sive volcanism, (Dietrich, 1979; Spray et al.,

1984), although individual outcrops are not large

enough to be shown on Plate 7. Large areas of

northern Europe were affected by Triassic exten¬

sion, controlling the subsidence of many basins

(Plate 7). Generally low Triassic sea levels and

narrow rift basins led to highly restricted water

bodies and the deposition evaporites (e.g. Muschel-

kalk, Keuper of northern Europe). During the Tri¬

assic, numerous marine connections were

established between Tethys and the basins of north¬

ern Europe, which, since the Lower Triassic, were

separated from the Arctic Seas.

On the Apulian block, a distinctive platform-

basin palaeogeography was established by Triassic

rifting. Grabens were filled with terrigenous sedi¬

ments, grading upward into pelagic, cherty carbon¬

ates. Platforms localized shallow water carbonates

(including reefs) and evaporites. In northern Italy,

Middle Triassic to Carnian volcanics and massive

reefs outcrop in the Dolomites. In the Adriatic and

Central Apennine areas, shallow water evaporites

(e.g. Burano formation) were deposited, while

deep-water elastics (e.g. Riva di Solto formation)

characterized the Northern Apennines and Po

Basins. Both the Burano and Riva di Solto forma¬

tions include source-rock facies (Anelli et al., this

52

P. O. YILMAZ ET AL TECTONIC EVOLUTION AND PALEOGEOGRAPH Y

volume). During Late Triassic times, possible ini¬

tial opening of the Vardar ocean occurred. This

marked the first formation of post-Hercynian

oceanic crust in the study area (Spray et al., 1984;

Dietrich, 1979).

In northern Apulia, distinct transverse zones,

following Hercynian trends, were established.

These zones were manifested as a series of flexures

which delimited domains of platform/basin geome¬

try.

During Late Triassic times, long-standing

south-directed subduction of the Palaeo-Tethys

ocean along the northern margin of the Cimmeria

blocks terminated with their collision with Europe

(Sengor, 1984; Sengor et al., 1984). In Europe,

Cimmerian deformations of Late Triassic and

Early Jurassic age is documented by unconformi¬

ties in the Polish Trough, the Northern Po Basin,

and in the Italian Dolomites.

Along the African margin, Ladinian to Carn-

ian age extensional tectonics created half-grabens

and pull-apart basins in the High Atlas trough. This

trend followed an aborted Late Carboniferous to

Permian rift (Cousminer and Manspeizer, 1977).

During the Late Triassic and Early Jurassic, the

newly created grabens were filled with continental

and evaporitic sediments. Further west in Morocco,

grabens were associated with continental elastics

and alkaline volcanism (Wildi, 1983).

Early Jurassic (Toarcian, 179 Ma)

During Early Jurassic times the Tethyan rift

system remained active (Biju-Duval et al., 1977;

Dewey et al., 1973; Dercourt et al., 1986; Ziegler,

1988, 1990). Plate 8 shows palaeogeography for

the Toarcian stage of the Early Jurassic. This was

also the time of initiation of opening of the Central

Atlantic between North America and Africa; with

this a sinistral strike-slip regime was established

between Africa and Europe. At this time marine

connections were reopened between the Arctic

Seas and the Tethys Ocean via the Arctic-North

Atlantic rift (Ziegler, 1988). Continued tectonic

subsidence of the Tethyan and European rift sys¬

tems, combined with a eustatic sea level rise, open

oceanic circulation patterns and low palaeolati-

tudes favoured deposition of widespread carbonate

platforms, especially on Apulia. In the Brian?on-

nais area (#4 on Plate 8), rifting caused foundering

of the older Middle to Upper Triassic platforms, on

which shallow water carbonates had been deposit¬

ed (Rudkiewicz, 1988; Michard and Henry, 1988).

In the Helvetic realm, sedimentation was

dominated by carbonates grading into marly

sequences (Funk et al., 1987). In Lias and Dogger

times, sedimentation consisted of mainly shales

(Dauphinois facies). Rapid horizontal facies

changes and wide stratigraphic gaps characterize

this facies (Masson et al., 1980). Rifting in the

Eastern Mediterranean resulted in formation of

oceanic crust in the Antalya area (# 2 on Plate 8;

Robertson and Dixon, 1984; Yilmaz, 1984).

Along the Cimmerian collision zone (#1 on

Plate 8), flysch deposition occurred in Dobrudgea,

Crimea and Northern Turkey (Sengor, 1984). Cim¬

merian orogenic activity terminated in Early Juras¬

sic times, as evidenced by intrusion of Middle

Jurassic plutons on both sides of the suture. Effects

of this event were also felt in the Polish Trough,

and local inversion occurred in the Donets Trough.

The Iberian Meseta was an important source

of elastics in the Cantabrian Basin, Lisbon and

Cavalla basins and also the Duero Basin (Wildi,

1983; Ziegler, 1988). Scattered extensional mag-

matism occurred on the Iberian Meseta and also in

the Pyrenees area.

Along the African margin, tilting and founder¬

ing of fault-blocks occurred mainly during the

Sinemurian and thus is slightly older than the time

represented on Plate 8 (Favre and Stampfli, 1991).

Middle Jurassic (Bathonian, 158.5 Ma)

The Bathonian map (Plate 9) shows the onset

of sea floor spreading in the Central Atlantic. This

spreading system continued to the north between

Iberia and Africa and then bifurcated into two

branches, one between Apulia and Europe and one

between Apulian and Africa. Continuation of the

northern branch past the Tisza and the Pienniny

area into the Black Sea is speculative; this is based

on occurrence of rifting in the Magura Trough and

PERI-TETHYS MEMOIR 2: ALPINE BASINS AND FORELANDS

53

Mecsek Zone and provides a mechanism for the

formation of the Transylvanian ophiolites.

In the Helvetic domain, sedimentation patterns

indicate progressive starvation and deepening of

the basin: Middle Jurassic manganese oozes are

succeeded by Late Jurassic radiolarites, Tithonian

Calpionellid oozes, slump deposits and transported

turbiditic calcarenites. In the Tethys area, a global

high-stand led to widespread carbonate platform

development. Pelagic sediments were deposited in

troughs while shallow water carbonate sedimenta¬

tion appears to have kept pace with subsidence on

the platforms, continuing the distinctive basin/plat¬

form topography of Apulia. Carbonate platforms in

the Dinaric and Friuli areas supplied large amounts

of debris into the Belluno Basin (Massari et al.,

1983). Apulian carbonate platforms reached their

maximum extent in Cretaceous time.

Late Jurassic-Earliest Cretaceous

We do not have palaeogeographic maps cover¬

ing Late Jurassic and earliest Cretaceous times.

Some of the major tectonic events are briefly sum¬

marized here.

Late Jurassic to Early Cretaceous tectonic

development of southern Europe was primarily

controlled by opening of the Central Atlantic

which established a regional sinistral shear

between Africa and Europe. Apulia rotated in a

counter-clockwise direction, opening the Mediter¬

ranean until Mid-Cretaceous times, when it

became fixed to Africa (Dewey et al., 1973;

Bernoulli and Lemoine, 1980; Biju-Duval et al.,

1977; Channell et al., 1979). Several ophiolitic

suites (e.g. Liguride, Piedmont, Transylvanian and

Vardar) were formed by these spreading events.

A Late Jurassic subduction zone with calc-

alkaline volcanism and flysch sedimentation

formed along the Dinaric shelf margin of Apulia

(Frasheri et al., this volume). Faunal evidence from

the Mecsek and Villany-Bihor areas on the Tisza

platform and the Brian?onnais zone of the Alps

(Roux et al.. 1988) indicates that these areas were

linked until Tithonian times and but were separated

afterwards. This indicates decoupling of the Tisza

block in Tithonian time. We suggest that Tisza sep¬

arated from Europe in Tithonian times and subse¬

quently became attached to Apulia before again

colliding with Europe in Eocene times.

Lower Cretaceous (Aptian, 112 Ma)

By Mid-Cretaceous times, Atlantic sea floor

spreading had propagated northward between

Iberia and North America and Iberia started to sep¬

arate from Europe (Plate 10). Mediterranean sea

floor spreading was nearly complete. Motions

between Africa and Europe, which were sinistral

from Jurassic through Early Cretaceous times,

changed in Mid-Cretaceous times to progressively

more convergent, with the convergence direction

becoming almost normal to the European margin

by Eocene times as the Arctic-North Atlantic

opened. This convergence established the Alpine

orogeny as well as several other, more localized

compressional events. One of these was in the

Rhodope area, with congressional deformation

and flysch deposition along the southern margin of

the Moesian Platform. Compression also occurred

along the eastern margin of Golija and on the

Pelagonian Platform. During Early Cretaceous

times, Pelagonia collided with Rhodope and

emplacement of nappes took place in the Hel-

lenides and Dinarides.

In Iberia and the Aquitaine Basin, block tilting

associated with extensional tectonics as well as

halokinetic movements of Triassic evaporites took

place during Early Cretaceous times (Le Vot et al.,

this volume). In the Western Alps, collision started

during Cenomanian-Early Senonian time due to

subduction of the European margin south or south¬

east beneath the Apulian margin; this is evidenced

by blueschists and eclogites (Debelmas, 1989). The

suture is seen in the Canavese slices (schistes lus¬

tres) and Sesia zone. High pressure metamorphism

associated with the suturing event has been dated

at 130 Ma and also 100-80 Ma (Debelmas, 1989).

Upper Jurassic -Lower Cretaceous ophiolite-bear-

ing, highly deformed nappes are overlain by Upper

Cretaceous relatively undeformed flysch nappes,

indicating the beginning of European-African com¬

pression in the Albian. Intra-Apulian deformation

was localized along pre-existing lines of weakness,

54

P. O. YILMAZ ET AL.: TECTONIC EVOLUTION AND PALEOGEOGRAPHY

principally Permo-Triassic grabens. An east-dip-

ping subduction zone initiated in front of the Golija

and Pelagonian platforms. During the Late Creta¬

ceous, major tectonic movements affecting the

Austro-Alpine domain, the internal Dinarides.and

the Southern Alps are expressed by a transition

from flysch sedimentation in the Lombardian and

Julian-Slovenian basins to a pelagic setting in the

Belluno basin and Trento platform (Massari et al.,

1983).

In the Apuseni mountains, Albian thrusting

and folding, along with flysch sedimentation, indi¬

cates that the Tisza block collided with or was

close to Apulia by Middle Cretaceous times

(Burchfiel, 1980). By Late Cretaceous times, large

strike-slip movements dominate the Tisza block

and North Pannonian part of the Apulian block as

the two blocks impinged on the Carpathian embay-

ment. This deformation is also expressed in the

Eastern Alps by lateral extrusion structures

(Ratschbacher et al., 1991). Further to the east, a

north-dipping subduction zone, characterized by

magmatic activity and back-arc rifting, was estab¬

lished in the present Black Sea (Gorur, 1989).

Counter-clockwise rotation of Iberia relative

to Europe resulted in sinistral shear along the

North Pyrenean fault zone (Galdeano et al., 1989;

Choukroune et al., 1989; Roure et al., 1989).

Oceanic crust developed in the Bay of Biscay fol¬

lowing early Aptian separation between Galicia

Bank and Flemish Cap (Dewey et al., 1973;

Ziegler, 1988, 1990). Rifting movements decreased

considerably in the East-Iberian Basin during Cre¬

taceous time, with continental to deltaic sandstone

deposition. In addition to the Pyrenean area, shear

deformation took place in the Cantabrian Moun¬

tains of northern Spain and in the Celt-Iberian

Range in central Spain. The nature of this shear

was predominantly transtensional and was mani¬

fested in the form of rifts and pull-apart basins.

Post-rift deposition began during late Aptian-

Albian time. A thick sequence of shallow water,

interbedded elastics and carbonates accumulated

on the subsiding Atlantic margin.

The Late Cretaceous was characterized by the

same opposed evolution of a shallow carbonate

platform in the eastern Iberides and a deep, terrige¬

nous flysch basin in the western Pyrenees. During

Mesozoic times, both on the platforms and in the

basins, the depositional sequence organization was

closely linked to eustatic sea level changes. Local

extensional processes generated important modifi¬

cations in thickness, particularly within the Lower

Triassic, Liassic, Kimmeridgian and Aptian series,

as well as within the Pyrenean Mid- and Upper

Cretaceous deposits. By the early Campanian, sea

floor spreading ceased in the Bay of Biscay and

Iberia began to converge with Europe. In the east¬

ern Pyrenees, the main deformation was Santonian

and Campanian. The Pyrenean collision front prop¬

agated westward during late Senonian to

Palaeocene time.

CENOZOIC PALAEOGEOGRAPHY

Lower Oligocene (Rupelian, 33.5 Ma)

Plate 1 1 shows palaeogeography for lower

Oligocene times. This was a period of intense tec¬

tonic activity following the collision of the Apulian

and European blocks. This collision was the result

of continued convergence between Africa and

Europe. The Apulia-Europe collision was diachro¬

nous, starting north of the present-day Adriatic and

propagating eastward into the Carpathians and

westward toward the Western Alps. Extensive

deformation, metamorphism, plutonic activity and

deposition of thick flysch and molasse sequences

occurred along the entire deformation front.

In the Alps, initial deformation occurred in the

Piemont, Briangonnais and Valais zones with later

involvement of the Ultrahelvetic and locally the

Helvetic domains. Development of the Molasse

Basin accompanied this deformation phase

(Ziegler et al., this volume). Northward transgres¬

sion of the Tethys sea led to a progressive onlap of

Cenozoic rocks onto the basal Tertiary unconfor¬

mity (Roeder and Bachmann, this volume). Local

positive features in the foreland persisted until

Oligocene time when the basin deepened rapidly

(Bachmann et al, 1987). Rising sea levels and local

restricted circulation provided ideal conditions for

the accumulation of very rich source-rocks (e.g.

Fish Shales formation in the Molasse Basin).

Source :

PERI-TETHYS MEMOIR 2: ALPINE BASINS AND FORELANDS

55

Alpine deformation was coeval with orogenic

activity along the Apennine and Dinaric fronts. Ini¬

tial deformation of the palaeo-Apennine chain

started during Oligocene times with thrusting of

the Liguride oceanic and flysch units (#2 on

Plate 11). Deformation and flysch sedimentation

were controlled by the palaeotectonic framework

inherited from Mesozoic tectonics. The shape and

interrelations of different basins varied as the struc¬

tural framework became better defined. The Apen¬

nine depocenters shifted from south to north as

different structural units deformed. The amount of

deformation increases from the north to the south

(Bally et al, 1986). This areal distribution of defor¬

mation controls trap size and trap integrity (Anelli

et al., this volume).

During Eo-Oligocene time, the Apulian

promontory pushed the North Pannonian and Tisza

blocks into the Pannonian embayment (Plate 11).

Together they “escaped” into this embayment,

which existed as a gap between the buttresses of

the Bohemian Massif to the northwest and the

Moesian platform to the southeast. In this process,

the North Pannonian and Tisza blocks pushed the

flysch, which had been deposited in front of the

inner, crystalline part of the Carpathians, over the

Carpathian foreland (#1 on Plate 11). The flysch

was folded and thrusted, accommodating at least

200 km of shortening. Loading of the foreland and

a coincident high stand in sea level provided ideal

conditions for the accumulation of widespread,

very rich Oligocene source-rocks in the region of

the Carpathian fold-and-thrust belt (e.g. Dysodilic

and Menilitic shales; see Bessereau et al., Dicea,

this volume).

Compressional deformation of the Alps and

their foreland was contemporaneous with the evo¬

lution of the intracratonic Rhine, Bresse, and

Rhone rifts (Ziegler, 1987, 1990; Ziegler and

Roure, this volume). A pulse of volcanism may

have triggered extension in the Rhine Graben dur¬

ing the late Eocene (#4 on Plate 11). By late

Eocene-early Oligocene times, a marine connec¬

tion was established between the Alpine foredeep

and the North Sea via the Rhine Graben in which

the Fish Shales formation was deposited, the

source-rock for most of the hydrocarbon accumula¬

tions in this graben.

Continued convergence of Iberia with Europe

caused westward propagation of the Pyrenean

deformation front into the Cantabrian region. Dur¬

ing the Eocene main Pyrenean deformation phase

Iberia was sutured to Europe (Roure et al., 1989;

Chouckroune et al., 1989; Ziegler, 1988, 1990).

Tethys ocean crust subduction continued in

the Alboran-Corsica/Sardinia region (#5 on

Plate 11). Back-arc rifting behind this northwest¬

dipping subduction zone separated the Balearic

Islands, Kabyl, Alboran and Corsica/Sardinia

blocks from Europe (Tome et al., Vially and Tre-

molieres, this volume).

Middle Miocene (Serravallian, 10.5 Ma)

Miocene Alpine deformation strongly affected

several parts of the orogenic belt (e.g. Central,

Western and Southern Alps, Carpathians, Apen¬

nines, Plate 12). In the Molasse Basin, with falling

sea levels and increasing influx of elastics from the

rapidly advancing Alpine orogenic front, the basin

shallowed and a thick continental molasse section

was deposited (Roeder and Bachmann, this vol¬

ume).

During Middle Miocene time, back-thrusting

of the Southern Alps accompanied deformation

along the Dolomites (Doglioni, 1991, 1992;

Ziegler et al., this volume). These thrusts involved

the Apulian Mesozoic platform and raised the

northern Po area (#3 on Plate 12). In the southern

Po area, local emergence and evaporite deposition

(e.g. Gessosso Solfifera formation) occurred.

Active west-dipping subduction in the Apennine

region and development of an east-facing foredeep

took place, recorded by the Macigno flysch (Anelli

et al., this volume). Flysch sedimentation from

both the Apennine and Albanian accretionary com¬

plexes created the Po/ Adriatic basin as a bivergent

foredeep area. The main thrusting event in the

Apennines occurred during the Miocene. These

thrusts utilized Triassic evaporite layers as detach¬

ment surfaces (D'Argenio et al., 1980, Boccaletti

and Coli, 1982). Tortonian continental and lacus¬

trine sediments, unconformably overlying flysch

facies, record this event, thereby constraining the

timing of the end of the main deformation phase.

On the east side of the Adriatic, subduction of the

relict Tethyan ocean continued along the Albanian

56

P. O. YILMAZ ET AL.: TECTONIC EVOLUTION AND PALEOGEOGRAPH Y

and Hellenic subduction boundary (#2 on

Plate 12).

The Carpathian foredeep expanded over the

European margin as the deformation front migrated

to the East- and South-Carpathians (#4 on

Plate 12). Rapid advance of the thrust front caused

the Oligocene source section to be uplifted and

maturation of the rich source facies terminated.

Therefore, over large areas, Oligocene source-

rocks are only mature where they have been struc¬

turally buried in the thrust belt or buried by

sediments in the very proximal parts of the fore¬

deep (Bessereau et al., Ziegler and Roure, this vol¬

ume). Volcanism around the Carpathian arc, related

to this phase of deformation, began during the late

Oligocene and is thought to be related to the sub¬

duction of highly attenuated European continental

crust beneath the overriding Carpathians. Back-arc

extension began in the Pannonian basin during this

time. By the end of the Miocene, deformation had

stopped in all but the Romanian portion of the

Carpathians (#5 on Plate 12).

By early Miocene time, the marine connec¬

tions through the Rhine, Leine, and Eger grabens

were severed as a result of uplift of the Rhenish

Massif and Massif Central. Tectonism in these

grabens, in which as much as 3 km of continental

sediments were deposited, continued, as shown by

volcanic activity.

Corsica/Sardinia and the Kabyl blocks were

separated from Europe during the early Miocene

(23 to 19 Ma; #1 on Plate 12; Vially and Tre-

molieres, this volume). This motion was a primary

driver for deformation of the Apennines. Fault

blocks formed by rifting in the Valencia Trough

contain the largest oil play found to date in Spain

(Torne et al., this volume).

The Late Miocene was characterized by com¬

pression along the southeastern margin of Iberia in

the Prebetic fold belt and in portions of the

Balearic Islands. Westward escape of the Alboran

Block opened the North Algerian Basin in its

wake. The Alboran Block collided with the south¬

eastern margin of Iberia and the northwestern mar¬

gin of Africa during the late Oligocene/early

Miocene. Extensive flysch basins mark this

Betic/Rif deformation front (the Numidian flysch

of Wildi, 1983; Ziegler, 1988). The Kabyl block

escaped southwards and collided with North Africa

to form the Tellian mountains of Algeria and

Tunisia (Wildi, 1983).

During the late Tortonian, Calabria rifted from

the Corsica and Sardinia block, creating the

Tyrrhenian Sea as a back-arc rift basin.

Lower Pliocene (3.8 Ma)

During Pliocene time, development of ocean

crust in the Tyrrhenian Sea was accompanied by

extensive magmatism (Channell and Mareschal,

1989, #1 on Plate 13). This was synchronous with

Pliocene Apennine nappe emplacement. Shorten¬

ing in the Apennines (75 km in the north. 150 km

in the south; Bally et al, 1986) was balanced by

extension in the Tyrrhenian Sea (Doglioni, 1991).

Extension initiated in late Tortonian time in the

Apennine hinterland, creating small rift basins

filled with Neogene elastics, sitting piggyback-

style on the thrust belt (Boccaletti and Coli, 1982).

The Apennine foredeep shifted northwards as the

Liguride thrust belt was reactivated during the

Pliocene. The Pliocene section, 9 km thick, con¬

sists of shallow water fluviatile sediments (#4 on

Plate 13). Tertiary biogenic gas plays are found in

this foredeep (Anelli et al., this volume). The dra¬

matic subsidence in this foredeep can not be

explained alone by the topographic load of the

Apennine thrust sheets (Royden and Kamer, 1984;

Royden, 1993).

In the Alps, the most significant deformation

is in the Helvetic domain. The deformation front

continued to progress to the north, leading to late

Miocene and early Pliocene folding and thrusting

of the Jura Mountains. Peak deformation was dur¬

ing the latest Pliocene. The thrust decollement in

the Jura is located in the Triassic evaporite section

which continues to the south under the Molasse

Basin and eventually under the Alps. The Molasse

Basin was carried passively to the north (as a

“piggy-back" basin) during this part of its history

(Philippe et al., Ziegler et al., this volume). By the

end of Tertiary times an approximately 5 km thick

section of synorogenic elastics had accumulated in

this basin.

Deformation continued in the outermost East-

and South-Carpathians. Extremely rapid subsi-

Source :

PERI-TETHYS MEMOIR 2: ALPINE BASINS AND FORELANDS

57

dence in the East-Carpathian foredeep occurred

during Mio-PIiocene time with accumulation of

approximately 9 km of coarse clastic molasse sedi¬

ments. The East-Carpathian foredeep is intensely

deformed to the west where it is partially overrid¬

den by thrusts of the flysch zone. To the north, this

inner zone, which consists mainly of the lower

molasse, is deformed by thrusts and folds. Salt

appears to act as a detachment surface but is also

involved in folding as salt diapirs pierce some of

the folds and salt is locally squeezed up along

thrust faults. This section contains the giant fields

of the Ploesti district (in the area of SCF on

Plate 13; Dicea, this volume).

The dramatic subsidence in the East-Carpathi¬

an foredeep can not be explained by the modest

topographic load of the Carpathian Mountains

(Royden, 1993, Doglioni, 1992). Royden (1993)

suggests that the Carpathians are an example of a

retreating subduction boundary where overall plate

convergence is less than the rate of subduction.

The deficiency in plate convergence was compen¬

sated by extension in the Pannonian back-arc

basin. Regional extension continues today as seen

in high grade metamorphic rocks of mid-crustal

origin which are exposed in Rechnitz window and

in Hungary (Tari, 1991).

The Carpathians are tectonically active today

with activity largely restricted to the area of the

Carpathian bend, also known as the Vrancea seis¬

mic zone. Large earthquakes with very deep

hypocenters and compressional to strike-slip fault

plane solutions typify this area; these may be pro¬

duced by the leading edge of the (detached) down¬

going slab under the Carpathian arc.

CONCLUSIONS

In this study, we utilized the plate tectonic his¬

tory of Europe to constrain the understanding of

sedimentary basin development and the effects of

regional scale tectonic events on play elements for

major basins. The tectonic framework and palaeo-

geography were used as constraints on models for

basin formation, climate distribution and accom¬

modation space which, in turn, control the distribu¬

tion of reservoirs, source-rocks, seals and traps.

The structural and stratigraphic framework of

Europe is the result of its Phanerozoic tectonic his¬

tory, involving the amalgamation of crustal blocks.

Multiple rifting and collision events created

extremely complex mountain systems during the

Caledonian, Hercynian, Cimmerian and Alpine

orogenies. Basins are diverse, superimposed, have

long-lived tectonic histories with complex structur¬

ing, and have highly variable play elements. The

Hercynian orogen provides the framework for

North European hydrocarbon systems. Its collapse

sets up the Apulian Mesozoic hydrocarbon system.

Alpine deformation and tectonically related exten¬

sion, in turn, set up the Neogene hydrocarbon sys¬

tems of the Carpathians, Pannonian Basin and the

Apennines (Ziegler and Roure, this volume)

Acknowledgments - We thank Exxon Produc¬

tion Research Company (EPR) and Exxon Explo¬

ration Company (EEC) for permission to publish

this paper. Also, we express our special thanks to

EP R-Geoscience for funding of this publication.

Hoa Tran of EEC drafted all the maps. We wish to

thank the other members of the European Regional

Study Team for their contributions: R. S. Bishop,

H. M. Bolas, L. B. Cauffman, S. Gardiner, D.

Gilbert, M. H. Feeley, J. L. Hagmaier, P. M. Haller,

M. T. Ingram, S. D. Knapp, R. A. Kolarsky, D. W.

Mason, R. T. Mooney, J. A. Newhart, G. J. Nolet,

L. S. Smith and J. P. Verdier. This manuscript ben¬

efited from earlier reviews by Mike R. Hudec. We

thank Peter Ziegler for inviting us to contribute to

this volume.

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Enclosures

Enclosure 1 Paleozoic crustal blocks on Per¬

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Enclosure 2 Mid-Carboniferous Namurian

(322 Ma) paleogeography

Enclosure 3 Upper Carboniferous Westphalian

A/B (306 Ma) paleogeography

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(254 Ma) paleogeography

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Enclosure 7 Upper Triassic Rhaetian (210 Ma)

paleogeography

Enclosure 8 Lower Jurassic Toarcian (179 Ma)

paleogeography

Enclosure 9 Middle Jurassic Bathonian

(158.5 Ma) paleogeography

Enclosure 10 Lower Cretaceous Aptian (122 Ma)

paleogeography

Enclosure 1 1 Lower Oligocene Rupelian

(33.5 Ma) paleogeography

Enclosure 12 Middle Miocene Serravalian

(10.5 Ma) paleogeography

Enclosure 13 Lower Pliocene (3.8 Ma) paleo¬

geography

Source :

Accretion and extensional collapse

of the external Western Rif (Northern Morocco)

J. F. Flinch

Department of Geology and Geophysics,

Rice University, Houston,

TX 77251-1892, USA

Present address : Departamcnto de Geologia,

Lagoven, Caracas 1010-A-889, Venezuela

ABSTRACT

The frontal part of the Gibraltar Arc consists

of allochthonous tectono-sedimentary complexes

classically interpreted as gravity driven units or

“melange”. High quality seismic data along the

northwestern Moroccan Atlantic margin and the

Rharb Basin, as well as field data in the Western

Rif provide a new view of this complex region.

The overall type of deformation suggests an accre¬

tionary prism involving deep-water sediments, that

was emplaced on the attenuated passive margins of

Iberia and Africa in response to westward motion

of the Alboran domain during the Miocene. The

timing of deformation and the age of the sediments

involved suggest an accretionary progression

towards the external portion of the Arc.

Late Miocene and Pliocene extensional col¬

lapse of the unestable accretionary prism controls

the structure of the region. The geometry of the

extensional system present in the frontal accre¬

tionary wedge of the Rif Cordillera is very similar

to the one of the Gulf of Mexico.

Data presented here suggests that some units

classically interpreted as thrust sheets are in fact

mixed extensional-compressional “satellite”

basins.

1 REGIONAL SETTING

The Western Mediterranean region consists of

a collage of several blocks located between the

Euro-Asiatic and the African plates. These inter¬

mediate blocks or micro-plates (i.e. Iberia,

Alboran, Corsica-Sardinia and Apulia) interacted

with each other and with Eurasia and Africa, defin¬

ing the geodynamics of the region (Dercourt et al.,

1986; Andrieux et al., 1989; Dewey et al., 1989;

Ziegler, 1987; Favre and Stamfli, 1992).

The Gibraltar Arc is the western limit of the

Alpine-Mediterranean system. The Betic

Cordillera in southern Spain and the Rif Cordillera

in northern Morocco constitute the northern and

southern part of the Arc (Fig. 1). The geological

units of the Betic and Rif Cordilleras can be subdi¬

vided, as most of the Alpine orogens, into an

External and an Internal domain.

Flinch, J. F„ 1996. — Accretion and extensional collapse of the external Western Rif (Northern Morocco). In: Ziegler, P. A. «&

Horvath F (eds) Peri-Tethys Memoir 2: Structure and Prospects of Alpine Basins and Forelands. Mem. Mus. natn. Hist, not ., 170: 61-85

+ Enclosures 1-2. Paris ISBN: 2-85653-507-0.

This article includes 2 enclosures on a folded sheet.

Source :

modi Heel after Dercourt el al. (1986) and Dewey ct al. (1989).

62

J F. FLINCH: EXTERNAL WESTERN RIF. NORTH MOROCCO

c.

C-

£i

Source : MNHN, Paris

EURASIA

PERI-TETHYS MEMOIR 2: ALPINE BASINS AND FORELANDS

63

Classically, the External domain of an oro-

genic belt is represented by non-metamorphic sedi¬

mentary successions, characterized by thin-skinned

tectonics. The External domain of the Betic and

Rif Cordilleras is separated into a number of struc¬

tural units or thrust sheets (Garcfa-Hernandez et

al., 1980; Vera, 1983; Wildi, 1983). Some thrust

sheets consist primarily of platform carbonates and

other sheets consist mostly of siliciclastic sedi¬

ments. The different types exhibit widely different

styles of decollement tectonics. The External

domain of the Betic and Rif Cordillera represents,

respectively, the south-Iberian and north-African

passive margin successions incorporated into the

folded belt (Michard, 1976; Vera, 1981; Wildi,

1983; Martin-Algarra, 1987).

The Internal or Alboran domain, which

includes the Internal zones of the Betic and Ril

Cordilleras (Fig. 2), differs stratigraphically and

structurally from the External zones (Suter, 1965;

Fontbote, 1983). The most notable characteristics

that distinguish the Internal from the External

zones include: Alpine-type Triassic carbonates,

Early Alpine (Cretaceous-Paleogene) polyphase

compressional deformation and HP/LT metamor¬

phism (Fontbote, 1983; Wildi, 1983; Galindo-Zal-

divar et al., 1989; De Jong, 1991). The lowermost

unit, the Nevado-Filabrides, is only present in the

Betic Cordillera (Fig. 3a). The intermediate Alpu-

jarrides-Sebtides unit contains metamorphic rocks

and mantle peridotites (Beni-Bousera and Ronda

ultramafics). The upper unit (Ghomarides-

Malaguides) overlies both the Sebtides and the

Dorsale unit.

Neogene extension associated with the col¬

lapse of the Alboran Sea and late inversion and

transpressional tectonics severely modified the

overall compressional development of the Gibral¬

tar Arc (Garcia-Duenas et al., 1992; Flinch, 1993).

Equivalent structural units of the Betics and the Ril

are presently separated by the Alboran Sea. Exten-

sional delamination accounts for dramatical thin¬

ning of the Internal domain units, resulting in

stratigraphic omission (Garcia-Duenas and

Martmez-Martinez, 1989; Garcia-Duenas et al.,

1992). The structure of the extended Internal

domain is similar to core-complexes of the Basin

and Range province of the Western United States

(Galindo-Zaldivar et al., 1989).

Two generalized regional geological cross-

sections were constructed to compare the structure

of the Betic and the Rif Cordilleras (Fig. 3). These

sections arc in part based on reflection seismic data

(Blankenship, 1992; Flinch, 1993), well-logs

(IGME 1987) and surface geological data (Suter,

1980a, 1980b; Blankenship. 1992; Garcia-Duenas

et al., 1992; Flinch, 1993). They provide a broad

approximation of the Gibraltar Arc structure and

help to define the main structural problems. At a

first look both cross-sections show great similari¬

ties. The Rif and the Betic Cordilleras are both

characterized by foreland-vergent thin-skinned

piggy-back thrusting involving the passive margin

succession of the north-African or south-Iberian

domains and a main Triassic decollement. Exten-

sional structures affect the Internal domain of both

fold and thrust belts, often reactivating previous

thrust sheets (negative inversion). The Frontal

tectono-sedimentary unit, referred to as

Guadalquivir Allochthon in the Betic Cordillera

and the Prerifaine Nappe in the Rif, is equivalent to

both sides of the Gibraltar Straits.

The main difference between these fold belts

is that the Betic is a carbonate dominated margin,

while the Rif is a detritic dominated margin. The

principal problems of the area concern the role of

extension related to the opening of the Alboran

Sea, the role of strike-slip faults, the way per

which the shortening in the lower part of the pas¬

sive margin succession is accomodated, the area of

origin of the frontal allochthonous tectono-sedi¬

mentary complexes and the sequence of thrust

emplacement.

2 THE ACCRETIONARY ZONE

The frontal units of the Betic and Rif Cor¬

dilleras (i.e. Frontal tectono-sedimentary Complex)

consist of highly deformed Triassic, Cretaceous,

Paleogene and Neogene strata, which were

detached from their original base and thrust over

the Mesozoic to Lower Miocene of the foreland.

This unit is known as the “Guadalquivir allochtho¬

nous units” in the Betic Cordillera which is equiva-

Source :

64

J. F. FLINCH: EXTERNAL WESTERN RIF, NORTH MOROCCO

FIG. 2. Tectonic map of the Gibraltar Arc, after Flinch (1993). Location of the

study area and following figures.

Source : MNHN, Paris

SW RIF CORDILLERA

PERI-TETHYS MEMOIR 2: ALPINE BASINS AND FORELANDS

65

Source : MNHN, Paris

Blankenship (1992) and Garcia-Duenas et al. (1992).

66

J. F. FLINCH: EXTERNAL WESTERN RIF, NORTH MOROCCO

lent to the Prerifaine Nappe of the Moroccan Rif.

Both the Guadalquivir Allochthon and the Preri¬

faine Nappe constitute an up to 100 km wide belt

(Fig. 2). The frontal allochthonous units of the

Betic Cordillera (i.e. the Guadalquivir allochtho¬

nous units) extend into the Gulf of Cadiz and far¬

ther west to the so-called Horseshoe or “Fer du

Cheval”, east from Gorringe Bank and the Seine

abyssal plain in the Central Atlantic (Fig. 1 ) (Lajat

et al., 1975; Malod and Didon. 1975; Malod and

Mougenot, 1979).

Near Gibraltar the inner units of the Arc are

bounded by the so called Flysch units, which con¬

sists of deep-water shales and turbiditic sand¬

stones.

The Guadalquivir Allochthon incorporates

voluminous Triassic shales and evaporites, as well

as marls, carbonates and siliciclastics (Perconig,

1960-62). In contrast, the Prerifaine Nappe

includes only minor amounts of Triassic evapor¬

ites, but mainly younger detritic turbidites or shales

(Daguin, 1927; Termier, 1936: Bruderer and Levy,

1954; Tilloy, 1955a, 1955b, 1955c, 1955d). In the

following I will focus on the southern part of the

Gibraltar Arc, in the accretionary zone of the Rif

Cordillera, classically refered to as the Prerifaine

Zone (Suter, 1965).

2.1 Stratigraphy

The sedimentary prism of the accretionary

wedge can be subdivided into three major tectono-

stratigraphic units: Supra-Nappe, Nappe and Infra-

Nappe. Due to the lack of offshore wells (only a

shallow well was drilled offshore Larache, see

Fig. 6 for location) stratigraphic information off¬

shore is largely based on correlation with onshore

data (Flinch, 1993). Onshore wells were tied to

seismic sections and correlated with offshore data.

Section of Figure 5 illustrates an example of

onshore-offshore correlation.

In the study area few wells have penetrated

the sedimentary succession located below the Pre¬

rifaine Nappe (Fig. 5). Cretaceous and Lower to

Middle Miocene strata unconformably overlie

metamorphic and igneous Paleozoic rocks of the

Hercynian Basement. Locally a Triassic shaly and

evaporitic section was encountered. Seismic data

demonstrates that Triassic sediments occupy exten-

sional half-grabens. The Cretaceous and Lower-

Middle Miocene section represents the cover of the

Morocan Meseta. The lack of Jurassic in this area

is explained by some authors as a result of shoul¬

der uplift related to the opening of the Central

Atlantic (Favre et al., 1991; Favre and Stampfli,

1992). The sedimentary succession encountered by

exploratory wells in the Rharb Basin is similar to

the stratigraphic section exposed in the western

Moroccan Meseta described by Gigout (1951).

The Prerifaine Nappe consists of Triassic to

Miocene sediments that are bounded by strati¬

graphic and/or tectonic contacts (Bruderer and

Levy, 1954). The stratigraphy for the Prerifaine

Nappe is obscured by complex deformation.

Resedimentation and the presence of reworked

Cretaceous faunas (i.e. ammonites) together with

Triassic and Miocene sediments leads to difficult

biostratigraphic problems (Feinberg, 1986). Thick¬

ness estimates are not easily made. The stratigra¬

phy of the Prerifaine Nappe is based on very few

outcrops located in the areas surrounding the

Rharb Basin and on exploration wells located with¬

in the basin. The following description of the

stratigraphy is based on surface data obtained by

the SCP (Tilloy, 1955a, 1955b, 1955c, 1955d;

Feinberg, 1986) and information from exploration

wells that penetrated the Nappe. Table 1 describes

the lithology and fossil content of the Prerifaine

Nappe. The Neogene planktonic biozonation is

derived from Feinberg (1986) and Wernli (1988).

Even though there is not a discernable stratigraphic

order because of imbrication and reworking, the

sediments involved in the Nappe proceed from

older to younger.

The Supra-Nappe complex consists of a sea¬

ward prograding wedge that ranges in age from

Late Miocene to Holocene. However the presence

of restricted anoxic environments poor in plankton¬

ic faunas, resedimentation and tectonic complica¬

tions hinder the establishment of a generalized

biostratigraphy. The litoral faunas of the peripheral

regions of the Rharb Basin are difficult to correlate

with the pelagic faunas of the center of the basin.

The Neogene biostratigraphy of the Rharb and

Rif areas, was based on Wernli (1988). The Supra-

Nappe subsurface stratigraphy of the Rharb Basin

based on selected wells is shown in Fig. 5.

Source : MNHN. Paris

PERI-TETHYS MEMOIR 2: ALPINE BASINS AND FORELANDS

67

Source : MNHN, Paris

TABLE

FIG. 4. Cross-section of the Accretionary Wedge in the Western Rif outlining the Offshore-Onshore correlation.

Notice that the apparent gentle westward deepening of the wedge is a velocity effect due to the water column.

J. F. FLINCH: EXTERNAL WESTERN RIF. NORTH MOROCCO

TWT (SEC)

TWT (SEC)

Source : MNHN. Pahs

PERI-TETHVS MEMOIR 2: ALPINE BASINS AND FORELANDS

69

3 STRUCTURE OF THE EXTERNAL

WESTERN RIF

Reflection seismic, well-log and field data

were integrated to provide an understandable pic-

lure of the external part of the Western Rif. Subsur¬

face data is specially important in this area due to

the lack of good surface exposures. A large number

of seismic profiles covering the northwestern

Atlantic margin of Morocco were used to map the

structure of the Frerifaine Nappe in the offshore

region. Mapping permits to establish the offshore

prolongation of the Rif frontal thrusts, the leading

edge of the frontal accretionary wedge and the con¬

tact between extensional and compressional

provinces (see Fig. 6). A complex set of basins can

be outlined within the accretionary zone (Flinch

and Bally, 1991).

3.1 Regional Transects

Five NE-SW offshore regional sections,

extending from Asilah to Rabat, have been selected

to show the structure of the northwestern Moroc¬

can Atlantic margin (Fig. 6 and Enclosure 1). The

transects display a variety of structural styles.

According to the seismic character and structural

significance, a number of structural units are dif¬

ferentiated, from bottom to top:

(1) Acoustic basement: Paleozoic.

(2) Infra-Nappe: Mesozoic and Lower

Miocene cover of the Paleozoic basement.

(3) Prerifaine Nappe: Accretionary Wedge.

(4) Supra-Nappe: Upper Miocene and Plio-

Pleistocene siliciclastics.

In the following these units will be referred to

as: Basement, Infra-Nappe, Nappe and Supra-

Nappe. The NE-SW oriented sections (Enclo¬

sure 1) traverse the main structural units of the

accretionary complex shown on the structural map

(Fig. 6). The regional sections will be described

proceeding from the southern foreland basin to the

northern frontal folded belt. The southern part of

the transects shows northward-dipping layered

reflectors of the foreland which project under the

frontal imbricates of the accretionary complex.

These reflectors are occasionally detached from the

acoustic basement to form imbricates. The frontal

imbricates are characterized by thrust planes which

dip steeply to the north. Thrust sheets emanate

from a gently northward-dipping basal decollement

which separates them from the underlying

autochthon. Proceeding northward, the complex is

overprinted by northward-dipping normal faults

and associated extensional basins with no signifi¬

cant growth. Further north, extensional faults step

down and confine thick extensional basins. They

constitute the southern part of an extensional sys¬

tem running nearly perpendicular to the plane of

the sections. Ridges cored by folded accretionary

complex sediments occur in the central region of

the extensional system. The southern branch of the

extensional system is characterized by northward¬

dipping normal faults. These faults are connected

with conjugate southward-dipping listric normal

faults that often constitute the lateral ramps of the

extensional system. The central portion of the

extensional basin, with its high ridges and deep

troughs, is detached from the basal extensional

contact. This portion of the margin provides excep¬

tional sections across the extensional system. Pro¬

ceeding northward, the northern branch of the

extensional system cuts thrusts and folds of the

underlying accretionary complex. The northern

portion of the sections is characterized by north¬

dipping thrust planes and related ramp anticlines.

According to the data presented here, the

study region can be subdivided into several struc¬

tural domains (Fig. 6): Offshore Tanger-Asilah

Congressional Belt, Offshore Larache Extensional

Zone, Offshore Rharb Compressional-Extensional

Zone, Rharb Basin and Rabat Foreland Basin. To

facilitate the presentation of the data, the offshore

data will be presented separate from the data of the

Rharb Basin, despite their structural affinity.

FIG. 5. Stratigraphic correlation of some selected wells of the Rharh Basin. Data from Wernli (1988), Feinberg

(1986) and unpublished ONAREP reports.

70

J. F. FLINCH: EXTERNAL WESTERN RIF, NORTH MOROCCO

Source : MNHN, Paris

PERI-TETHYS MEMOIR 2: ALPINE BASINS AND FORELANDS

71

Offshore Tanger-Asilah

Fold and Thrust Belt

Asilah

Offshore Larache -

Extensional System

Larache

Rhabet Jeblla

Offshore Rharb

Extensional / Compressions!

Rharb Basin

Extensional / Compressions! zone

R7

Lai I a Yto

LEGEND

Kenitra

Foreland Basin

FIG. 6. Structural map of Northwestern Morocco with indication of the major

structural domains consider in this study.

Source : MNHN. Paris

72

J. F. FLINCH: EXTERNAL WESTERN RIF. NORTH MOROCCO

3.1.1 Offshore Tanger-Asilah / Fold and Thrust

Belt

The northern part of the Moroccan Atlantic

margin consists of westerly-vergent folds and

thrusts. Fold axes and thrusts strike NNW-SSE,

following the general trend of the Rif Cordillera

(Fig. 6). They represent the continuation of the

structures of the Tanger and Habt units exposed in

the western Tanger peninsula (Suter, 1980b). The

quality of the seismic data in this area is mediocre,

and there are some regions with virtually no useful

data.

3.1.2 Offshore Larache / Extensionul Zone

The frontal thrusts of the Western Rif

Cordillera extend further west into the region off¬

shore Tanger- Asilah. Thrusts and related folds are

cross-cut by NW-SE trending SW-dipping low-

angle listric normal faults (Enclosure 1). The struc¬

ture of this region is defined by troughs and ridges.

Extensional basins are bounded by several anosto-

mosing listric normal faults that sole out into a

basal low-angle detachment which offsets the top

of the accretionary wedge (Fig. 6). This network of

anostomosing faults results in extensional horses

that merge with each other and are superimposed

on the accretionary wedge. Normal faults trend N-

S, nearly parallel to the present day shoreline, and

dip towards the west. E-W oriented sections (see

Enclosure 1 ) show- strongly rotated blocks on the

hangingwall of the low-angle extensional detach¬

ment. In the eastern portion of these sections,

growth-faulting results in large Supra-Nappe

expansion controlled by westward-dipping normal

faults that sole out into the basal detachment. Sedi¬

ments above the extensional system show signifi¬

cant fault growth. In the central portion of the area,

the top of the Nappe attains depths of 3.5 sec

(TWT). The basal detachment of the extensional

system steps down from 0.5 sec in the east to

3.5 sec in the west. The aparenl transport direction

of this extensional system is to the west. Nearly E-

W oriented shale ridges are present in the central

part of the extensional system; these were caused

by shale withdrawal induced by extensional dis¬

placement.

3.1.3 Offshore Rharb / Frontal Imbricates -

Extensional-Compressional Zone

This area is located west of the southern

Rharb Basin, between the confluence of the Sebou

River and the village of Moulay Bou Selham. High

quality seismic data in this region show the details

of the frontal part of the accretionary complex and

the northernmost portion of the Rif foredeep. The

Offshore Rharb area displays a combination of

compressional and extensional elements (Fig. 6).

In this complex area, NW-SE trending normal

faults and occasional NE-SW-oriented normal

faults cut NW-SE trending SW-vergent folds and

thrusts. To the east of the area, normal faults share

the same decollement level as thrust faults, defin¬

ing toe-thrusts that accommodate the normal fault

displacement (see Enclosure 1).

The frontal part of the wedge is characterized

by closely spaced NW-SE trending, NE-dipping

thrust faults that define a zone of frontal imbri¬

cates. Lateral and oblique ramps related to these

frontal thrusts are common in the central and west¬

ern portions of the area (Enclosure I). The front of

the accretionary complex has a NW-SE orientation

(Fig. 6).

The structure of this area is well constrained

by high quality seismic data. Dip lines are those

trending perpendicular to the leading edge of the

accretionary complex, that is NE-SW. Strike lines

are those trending roughly parallel to the front of

the wedge, that is NW-SE.

The most conspicuous features shown by the

dip NE-SW sections are:

(a) The wedge-like geometry of the Prerifaine

Nappe.

(b) Imbrications within the Infra-Nappe

authochthonous succession.

(c) Frontal thrusts within the Nappe.

(d) Extensional faults crosscutting the Nappe.

(e) A northwestward-dipping basement

beneath the Nappe.

Source : MNHN. Paris

PERI-TETHYS MEMOIR 2: ALPINE BASINS AND FORELANDS

73

(0 The northernmost portion of the foredeep,

(g) Frontal slumps

The strike sections show the following struc¬

tural features:

(a) Steeply-dipping basement-involved faults

offsetting the base of the Infra-Nappe

units.

(b) Eastward thinning of the Prerifaine-Nappe.

(c) Lateral ramps of the frontal imbricates.

(d) Westward progradation of the Supra-

Nappc succession.

(e) Frontal slumps and detached units.

3.1. 4 Rharb Basin / Frontal Imbricates -

Extensional-Compressional Zone

The topographic Rharb Basin overlaps the

western front of the Rif Cordillera and its foreland

(Fig. 2). It is bounded to the east and north by the

frontal ranges of the Rif Cordillera and to the west

by the Atlantic coast. The southern limit is the

Paleozoic Moroccan Meseta. The surface expres¬

sion of the Rharb Basin is a fluvial-alluvial coastal

plain drained by the Sebou River. Subsurface data

presented in this paper display the structure of the

Neogene sedimentary succession of the Rharb

Basin and the underlying Prerifaine Nappe.

South of the imbricated zone represented by

the Asilah, Habt and Tanger units (Suter, 1980b)

most of the structure of the accretionary wedge is

controlled by extensional structures (see Fig. 6 and

Enclosures 1 and 2). Extensional basins are located

between the Rif fold and thrust bell and the leading

edge of the accretionary wedge (Prerifaine Nappe).

The structure of the Rharb Basin does not consist

of well defined half-grabens but it is composed by

a complex set of extensional and locally congres¬

sional basins (Fig. 6, Enclosure 2). These “satel-

lite>’1 basins are bounded by several anostomosing

listric normal faults which sole out into a basal

low-angle detachment offseting the top of the

accretionary wedge. Often extensional structures in

the rear are coeval with conpressional toe-thrusts at

the front of these “satellite" basins. The orientation

of main faults is variable and random. Eventhough

there is controversial data on the timing of devel¬

opment of these basins, most of them are Torton-

ian-Messinian in age. Growth is limited and most

of the supra-nappe sediments are characterized by

parallel bedding, which suggests fast extensional

collapse. Basal re-deposited shallow-water sand¬

stones and anoxic marls with occasional siltstone

beds fill these extensional basins. Anoxia is the

result of the extensional topography at the top of

the Prerifaine Nappe, which involves highs and

lows (Cirac and Peypouquet, 1983). In the cover

sequence, rapid facies changes through time sug¬

gest also rapid extensional collapse of the Nappe

(Flinch, 1993).

In the southern part of the Rharb Basin com-

pressional structures associated with the front of

the Prerifaine Nappe are observed (Enclosure 2).

They consist of NE-SW trending anticlines and

synclines related to SW-vergent imbricates of the

Prerifaine Nappe. The southern area is occupied by

E-W and N-S trending troughs associated with

extensional faults superimposed on top of the

Nappe. The structure of the central part of the

Rharb Basin consists of E-W, N-S and NE-SW

trending extensional troughs bounded by low-angle

listric normal faults (Fig. 6). Depocenters in excess

of 4500 meters occur southwest from Mechra bel

Ksiri. (Dakki, 1992). Three of these basins are

located in the contact between the foothills of the

Rif and the Rharb Basin, namely: the Lalla-Zhara,

Souk el Arba (Fig. 7) and Nouriat Satellite Basins.

All these basins are characterized by rearward

extension and frontal compression, which defines

large scale toe-thrusts. The sedimentary fill of

these basins, consists mostly of Tortonian-Messin-

ian marls, with occassional interbedded sandstones

and siltstones. Shallow biogenic gas, and locally

oil has been recovered from the Ain Hamra area

(see Fig. 7 for location). The presence of Creta¬

ceous source oil at a very shallow level, few hun¬

dred meters, suggests a connection between the

1 The term satellite basin is preferred to piggy-back basin because of the presence of normal faults combined or not with thrust faults. The

mechanics of these basins is therefore different than conventional thrust-related piggy-back basins in the sense of Ori and Friend (1984).

74

J. F. FLINCH: EXTERNAL WESTERN RIF. NORTH MOROCCO

LEGEND

SUPRA-NAPPE F

Torlonian

PRERIFAINE

NAPPE ^

| Eocene

J Cretaceous

Triassic

V=H

1 Km

FIG. 7. Structural map and cross-scction along the Souk el Arba Satellite Basin.

Source : MNHN, Paris

PERI-TETHYS MEMOIR 2: ALPINE BASINS AND FORELANDS

75

extensional decollement and the underlying thrust

faults within the Nappe.

The structure of the Rharb Basin is illustrated

by a series of composite regional seismic lines.

Enclosure 2 shows four regional dip sections and

eight strike lines. The dip lines are roughly orient¬

ed NE-SW, trending nearly perpendicular to the

leading edge of the Prerifaine Nappe. The strike

lines are oriented perpendicular to the transport

direction of the Nappe and parallel to its leading

edge. Enclosure 2 displays five northeastward

trending sections in the Lalla Yto area. Three E-W

oriented sections, extending from the Nouriat

region in the east to the Sebou region in the west;

four NE-SW trending sections; two in Lalla Yto in

the East, one through the central Rharb extending

from Rhabet Jebila in the north to Lalla Yto in the

south and one along the Sebou coastal area. On the

basis of seismic character and structural signifi¬

cance, the same seismic units are recognized as in

the offshore sections (Enclosure 1), that is: Base¬

ment, Infra-Nappe, Nappe and Supra-Nappe.

The most characteristic features evidenced on

the dip lines (odd numbered sections of Enclo¬

sure 2) are:

(a) The wedge-like character of the Prerifaine

Nappe.

(b) The northward dip of the basement and the

infra-nappe succession underneath the Pre¬

rifaine Nappe.

(c) Southward vergent imbricates involving

infra-nappe sediments.

(d) Extensional faults offsetting the top of the

nappe.

(e) Clinoformal patterns in the Supra-Nappe

units indicating southweslward prograda¬

tion.

The most conspicuous features evidenced in

the strike lines (even numbered sections of Enclo¬

sure 2) are:

(a) The thickness change of the Infra-Nappe

unit.

(b) Nearly vertical faults offsetting the Infra-

Nappe succession.

(c) Lateral ramps of the extensional system

(the oblique orientation of the section

reveals lateral ramps).

3. 1.5 Onshore-Offshore Rabat / Foredeep

South of the leading edge of the Prerifaine

Nappe, the structures consist of NE-dipping Infra-

Nappe units that plunges beneath the accretionary

wedge. Basement-involving nearly-vertical normal

faults disrupt the otherwise continuous Infra-

Nappe succession (see Enclosures 1 and 2). Some

of these faults account for a thickness change of

the infra-nappe succession but do not affect the

overlying sediments (Enclosure 2); these faults are

related to the Central Atlantic rift system (Flinch,

1993). Other faults do not account for any thick¬

ness change of the infra-nappe units but offset the

foreland succession, these faults are related to flex¬

ural extension induced by tectonic loading of the

foreland by the Prerifaine Nappe (Enclosure 1)

(Flinch, 1993) as seen also in other foreland basins

(Bradley and Kidd, 1991 ). An angular unconformi¬

ty which is onlapped by the foreland sequence cor¬

responds to the “basal foredeep unconformity” in

the sense of Bally (1989). Locally, sediments

derived from the craton, prograding into the fore¬

deep can be recognized in the southern part of the

area (see Enclosure 1). This region is located in the

offshore prolongation of the Rif foredeep, located

east from Rabat (see Figs. 2 and 6 for location).

4 DISCUSSION

The overall composition and type of deforma¬

tion of the external Betic and Rif domain suggests

the involvement of an accretionary prism consist¬

ing of deep-water sediments, that was emplaced on

the attenuated Iberian and African passive margins

in response to westward motion of the Alboran

domain during Middle Miocene to Pliocene time.

76

J. F. FLINCH: EXTERNAL WESTERN RIF. NORTH MOROCCO

The timing of deformation and the age of the sedi¬

ments involved suggest an accretionary progres¬

sion towards the external’ portion of the arc.

Deformation within the accretionary complex was

previously explained by several phases of deforma¬

tion and by gravitational tectonics (Feinberg, 1976:

Vidal, 1977; Feinberg, 1986). In contrast, a contin¬

uous accretion model, similar to current models of

more conventional accretionary complexes (e.g.

Dickinson and Seely, 1979; von Huene 1986),

appears to apply here. On the basis of field data

and seismic data, a block diagram of the accre¬

tionary complex was constructed (Fig. 8). Note

that the accretionary wedge is presumably under¬

lain by a normal to transitional continental crust

which dips towards the Mediterranean (A-type

subduction); this contrast with the more conven¬

tional accretionary wedges that are related to sub¬

ducting oceanic lithosphere (B-type subduction).

The style of extension present in the external

part of the Gibraltar Arc is similar to the Gulf of

Mexico (Worrall and Snelson, 1989). It consists of

listric normal faults rooted into a low-angle exten-

sional detachment composed of overpressured

shales and marls. In the offshore Larache area, the

transport direction of the extensional system coin¬

cides with the geometry of the continental slope,

and is nearly parallel to the present day shelf-

break.

There is no clear relationship between the

frontal congressional zone (Offshore Rabat) and

the rear extensional system described above. The

lack of seismic resolution and penetration in the

lower part of the seismic sections does not allow to

see if down-dip thrusts and folds share the same

decollement as the up-dip low-angle normal faults.

Therephore, it is difficult to demonstrate if exten¬

sional displacement is compensated by frontal

compression at a regional scale. The most frontal

parts of the Gibraltar Arc accretionary wedge may

represent congressional belts that are the result of

rear extension, as suggested by Platt (1986) for

other accretionary wedges. This type of deforma¬

tion would represent the response of the system to

the unstable oversteepened slope (critical taper the¬

ory) generated by the stacking of thrust slices with¬

in the wedge (Davis et al., 1983).

The three-dimensional diagram presented here

has some important implications for the geology of

the Belie and Rif Cordilleras. The Rif Cordillera

consists of numerous structural units, which classi¬

cally are referred to as “Nappes” (Suter, 1980b),

however many of these units do not have the attrib¬

utes of thrust sheets. Omission of strata or no

duplication of the stratigraphic section are common

to these units. A structural map of the Western Rif

was put together integrating offshore and onshore

subsurface data and field data based mostly on

published geologic maps (Flinch, 1993) (Fig. 9).

The relationship between the Prerifaine Nappe and

the underlying units of the External domain can

be observed particularly well in the Had-

Kourt/Teroual area (Fig. 10). Seismic data through

the area permits to see the relationship between

several stacked thrust-sheets (Fig. 11). The data

presented here suggest that several geologic units

referred to as “unites flottantes” or “Nappes

rifaines superieures” (upper thrust sheets) (Wildi,

1983), are in fact satellite basins which overlie the

accretionary wedge (Figs. 9, 10 and 11). These

basins were deformed at the same time as the

transport of the accretionary wedge towards the

foreland, in response to the collision of the

Alboran allochthonous terrane with the Iberian and

African foreland. In the past, some of these Satel¬

lite basins were interpreted as out-of-sequence

thrusting (Morley, 1992). Instead a model where

the structure is the result of the piggy-back

emplacement of sequences, that is, the upper units

were emplaced first, the later emplacement of the

lower units deformed the upper units from under¬

neath. The first unit to be emplaced was the accre¬

tionary wedge and the underlying more landward

passive margin units were emplaced afterwards

(Fig. 11). This lead to widespread structural envel¬

opment of the accretionary wedge (Flinch,

1993).The new concept presented here, significant¬

ly simplifies the structural framework of the Rif

Cordillera. I postulate that these units may have a

similar origin as the Neogene satellite basins of the

offshore and onshore Rharb area.

The Ouezzane Unit (Hottinger and Suter in

Durand-Delga, 1960-1962) of the external Western

Rif is the most obvious example of such a complex

of satellite basins. This unit was interpreted as a

thrust sheet or “nappe" located on top of the Preri¬

faine nappes (Suter, 1965: 1980b). In the so-called

flysch domain, the Numidian Unit represents the

highest structural unit which is located above the

underlying imbricates; also this unit may represent

Source : MNHN . Paris

Foredeep Frontal Imbricates Extensional System

Source : MNHN. Paris

J. F. FLINCH: EXTERNAL WESTERN RIF, NORTH MOROCCO

STRAITS OF GIBRALTAR

SEBTA

LEGEND

50 Km

FIG. 9. Structural map of the Western Rif integrating onshore and offshore data

alter Flinch (1993).

Source

PERI-TETHYS MEMOIR 2: ALPINE BASINS AND FORELANDS

79

LEGEND

Anticline

Extensions! V Throat fault A sync,,n*

fault ^

| Neogene and Quaternary | - ^

MeesJnlan

aandslonee

Oueziane unit

( Eocene to Wocene)

ZoumJ aandalone

( OUgocene )

[ " ] Prerlfalne Nappe

□ Upper Imbrlcatea

(Upper Cretaceoua to Eocene )

Lower ImOrlcalee

( Juraaatc-Neocomlan )

FIG. 10. Structural map of the Had Kourt-Teroual area. Modified from the Ser¬

vice Geologique du Maroc 1:50.000 scale maps of Had Kourt (1984) and Teroual

(1990).

Source : MNHN, Paris

Line drawings from seismic sections along the Had (Court -Teroual area.

80

J. F. FLINCH: EXTERNAL WESTERN RIF, NORTH MOROCCO

2

5

TWT(sec)

U ^ N) O

to N) — O

TWT(sec)

Source : MNHN. Paris

SE I WNW

PERI-TETHYS MEMOIR 2: ALPINE BASINS AND FORELANDS

81

a set of satellite basins. The Numidian sandstone

was deposited on top of imbricates involving pre¬

viously deposited turbiditic deposits of the Tanger

and Ketama Units. I suggest that the Numidian

Sandstone represents the turbiditic Satellite basin

fill which overlies the Oligocene-Early Miocene

accretionary wedge of the Gibraltar Arc. These

Satellite Basins appear to be detached from the

underlying units by thrust or normal faults, thus

simulating an independent thrust sheet. However,

unlike real stacked thrust sheets, younger sedi¬

ments overlie the deeper tectono-stratigraphic unit.

5 CONCLUSIONS

The frontal tectono-sedimentary complexes of

the Betic and Rif Cordilleras, the Guadalquivir

Allochthon and the Prerifaine Nappe, constitute an

accretionary wedge which was superposed on the

attenuated passive margin of Iberia and NW Africa

during the Miocene phases of the Alpine orogeny.

The structure of the accretionary wedge con¬

sists of frontal imbricates, ridges, toe-thrusts and

low-angle extensional detachments. The Supra-

Nappe sediments involve compressional-exten-

sional and extensional satellite basins trending

parallel and perpendicular to the Arc. Satellite

basins are not directly related to the opening of the

Alboran Sea; instead they are due to oversteepen¬

ing of the accretionary wedge and gravitational

gliding down the continental slope.

Very rapid extensional collapse affected the

accretionary wedge during Tortonian and Messin-

ian time. The paleogeographic evolution, involving

the superposition of deep-water facies onto shal¬

low-water sediments and the lack of significant

growth support this argument.

The style of extension is characterized by low-

angle listric normal faults, similar to the Gulf ol

Mexico. Extensional displacement is compensated

by frontal compression. Toe-thrusts are common

structures in the frontal part of the Gibraltar Arc.

An undisturbed upper part of the prograding

Supra-Nappe succession suggests that the accre¬

tionary wedge was stabilized during Pleistocene

time.

Acknowledgements - This study was financed

by a Fulbright fellowship provided by the Spanish

Ministry of Education and Science. I want to thank

the Office National des Recherches et d' Exploita¬

tions Petrolieres ( ONAREP ) of the Kingdom of

Morocco and PETROCANADA for providing the

data presented in this work. Special thanks go to

Mr. Bouchla, Mr. Morabet and Mr. Demnati for

their support. Helpful discussions on the geology

of the area were held with Mr. Zizi, Mr. Hcaine,

Mr. Dakki, Mr. Bouchlouck, Mr. Mahmoud and

Mr. Jobidon. Thanks are extended to Prof. A. W.

Bally for his help, guidance during the project and

suggestions. I am also grateful to Dr. Peter Ziegler

for his helpful review and excellent editorial work.

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APPENDIX

In the following I will describe the structure of

the seismic line drawings presented in Enclosure 2.

84

J F. FLINCH: EXTERNAL WESTERN RIF, NORTH MOROCCO

Regional Section R1

This N-S and NE-SW regional section extends

from Rhabet Jebila to Lalla Ylo. In the northern

portion of the transect a series of southward-dip-

ping growth faults offset the top of the Prerifaine

Nappe. The region of Rhabet Jebila-Lalla Zhara

consist of a northern extensional satellite basin and

a southern ridge. The basal decollement is located

at 2 seconds recording time. Thrusts connected

with this basal detachment are present within the

Nappe. The maximum thickness (2.5 sec)

(1800 meters) of the Supra-Nappe Neogene is

attained in the central part of the section. The

southern part of the section (Lalla Yto region) is

represented by lateral ramps merging into the basal

decollement of the Prerifaine Nappe. Antithetic

and synthetic normal faults also offset the Mio-

Pliocene boundary. Normal faulting is younger in

this region than in the northern Rharb. where the

Mio-Pliocene boundary is not offset.

Regional Section R3

This northeastward trending section follows

the Sebou River coastal plain and shows the

wedge-shaped Prerifaine Nappe. In the southeast¬

ern portion of the section, gently northcastward¬

dipping Infra-Nappe reflectors unconformably

overlie the Hercynian basement. The Supra-Nappe

prograding units onlap directly the basal foredeep

unconformity. Well EM-3, located in the foredeep

region just in front of the Nappe, penetrated the

whole sedimentary succession, encountering Creta¬

ceous and Tortonian-Messinian Infra-Nappe and

Plio-Pleistocene Supra-Nappe. North of the leading

edge of the wedge, the Infra-Nappe unit is charac¬

terized by a basal zone of imbricated layered

reflectors. The structure of the Prerifaine Nappe

itself is characterized by northeastward-dipping

thrust laults. Extensional faults are superimposed

on the Nappe. Supra-Nappe units show a south¬

ward prograding clinoformal pattern. Well MO-1

penetrated the Nappe and Infra Nappe units, reach¬

ing the Paleozoic basement.

Regional Section R5

This NE-SW oriented section located in Lalla

Yto illustrates essentially the same features as on

section R3. Again, the geometry of the wedge, the

basal imbricates of the Infra-Nappe unit, and the

normal faults that cut the Supra-Nappe succession

are the most interesting features shown on this pro¬

file. On the SE margin of the section, south of the

leading edge of the Prerifaine Nappe i.e. in the

foredeep, the Supra-Nappe units onlap directly on

the northeastward-dipping Infra-Nappe reflectors,

thus definnig the basal foredeep unconformity.

Well KC-1, located in front of the Nappe, reaches

the basement after penetrating Infra-Nappe Trias-

sic, Cretaceous and Middle Miocene.

Regional Section R7

This northeastward trending section is located

in the region of Lalla Yto, west of section R5. The

section shows the same features as section R7. In

the frontal part of the Nappe, extensional and com-

pressional structures are detached at the same

decollement level, thus defining toe-thrusts.

Regional Section R2

This section is oriented E-W, extending from

the region of Nouriat to the Atlantic Coast. The

line shows the subsurface expression of the contact

between the Rharb Basin and the frontal ranges of

the Rif. The eastern end of the section shows west-

ward-dipping listric normal faults, responsible for

the westward-thickening of the Supra-Nappe suc¬

cession. Most normal faults merge into a low-angle

extensional detachment. Rotated blocks develop in

the hangingwall of the extensional system. Exten¬

sion in the Nouriat area occurs mostly during

Messinian time. Proceeding westward, the struc¬

ture of the central Rharb Basin consists of a series

of westward-dipping listric normal faults. Lateral

ramps, suggesting a transport direction oblique to

Source : MNHN, Paris

PERI-TETH YS MEMOIR 2: ALPINE BASINS AND FORELANDS

85

the plane of the section, are common in the central

portion of the section. In the western part of the

section (Sebou region), the base of the Prerifaine

Nappe is at 3 sec recording time. Infra-Nappe lay¬

ered reflectors are well imaged. Occasional thrusts

are present within the Nappe. The Prerifaine Nappe

is not significantly affected by extensional faults in

this western area.

Regional Section R4

This E-W trending regional section is parallel

to R2 and extends from Nouriat to the Atlantic.

Most conspicuous on this section is the westward

deepening of the top-of-the-Nappe extensional

detachment. Ramps and flats define the geometry

of the basal extensional detachment. The exten¬

sional decollement deepens down to 2.5 sec in the

central part of the section. Supra-Nappe sediments

fill the downthrown depressions of the extensional

system. Only on the easternmost portion of the sec¬

tion is the Mio-Pliocene boundary offset by normal

faults. In the Sebou region the basal decollement of

the Prerifaine Nappe is imaged on the protile at

3 sec (TWT). Layered Infra-Nappe reflectors

underlie the basal decollement.

Regional Section R6 (see Fig. 4)

This section shows lisia and Africa, defining

the geodynamics of the region (Dercourt et al.;

1986; Andrieux et al., 1989; Dewey et al., 1989;

Ziegler, 1987; Favre and Stampfli, 1992).

Regional Sections R8, RIO, R12, R14

These sections located in the Lalla Yto area

are closely spaced (2 to 3 km). They are oriented

NW-SE, providing more examples of strike sec¬

tions across the Rharb Basin. Wells MA-101 and

MO-1 were tied into the seismic. The basal

decollement steps down from 1.8-2 to 3-3.2 sec

recording time. Thickness changes of the Infra-

Nappe succession coincide with the location of

nearly vertical faults that offset the basement. The

top of the basement and the overlying Infra-Nappe

succession dip to the northwest. The Prerifaine

Nappe is thrusted onto Cretaceous and Middle

Miocene Infra-Nappe sediments that overlie the

Hercynian basement of the Moroccan Meseta. Lat¬

eral ramps of listric normal faults offset the top ol

the Nappe and the overlying Supra-Nappe succes¬

sion.

Regional Section R16

This section across the central Lalla Yto area

is roughly oriented WNW-ESE. The section is

located near the leading edge of the Prerifaine

Nappe. A complete Infra-Nappe succession pene¬

trated by the KC-1 well consists of a Triassic half-

graben and is unconformably overlain (Post-Rift

unconformity) by parallel-bedded horizontal Creta¬

ceous and Miocene sediments. Plio-Pleistocene

units show a downlap pattern that suggests w'est-

northwestward progradation.

Enclosures

Enclosure 1 Line-drawings of regional seismic sections, offshore Northwestern Morocco.

Sections A, B, C, D, E: offshore Asilah-Rabat. Sections F, G, H, 1: offshore Larache.

Enclosure 2 Line-drawings of regional seismic sections: Rharb Basin, onshore Northwestern

Morocco.

Source : MNHN , Paris

Triassic- Jurassic extension and Alpine inversion

in Northern Morocco

M. Zizi

ONAREP, 34 avenue Al Fadila, Rabat, Morocco

ABSTRACT

The Early Mesozoic half-grabens of northern

Morocco form part of the regional extensional sys¬

tem which developed in conjunction with the open¬

ing of the Western Tethys. During the Tertiary,

Alpine collision of the African and European

plates, these half-graben system were inverted, giv¬

ing rise to the uplift of the High and the Middle

Atlas mountains. Similar, albeit less spectacular

inversion features occur in the Guercif area and in

the "Rides Prerifaines" of northern Morocco.

Reflection seismic data show that inversion of the

Guercif Basin involved the reactivations Mesozoic

basement faults. In contrast, the "Rides Preri¬

faines" correspond to an extensive detachment sys¬

tem which is decoupled from the basement at the

level of the Triassic evaporites. The geometry of

this complex system of Late Miocene-Pliocene

thrust faults and associated lateral ramps was pre¬

conditioned by the configuration of the Triassic-

Jurassic extensional faults.

Detailed structural and stratigraphic analyses,

combining surface geology and seismic data, great¬

ly advanced the understanding of the geological

history Northern Morocco and has led a reassess¬

ment of its hydrocarbon potential.

INTRODUCTION

The main structural elements of northern

Morocco are the Moroccan Meseta, the Rif fold

and thrust belt and the northeastern part of the

Middle Atlas Mountains (Fig. 1).

The Moroccan Meseta is upheld by the out¬

cropping. peneplained Hercynian basement and its

Mesozoic cover. To the North, the Meseta dips

gently under the Neogene fill of the Rif foreland

basin which is underlain by a thin Mesozoic cover.

This basin is subdivided into the Rharb Basin in

southwest, the south-central Saiss Basin and the

Guercif Basin in the northeast. These basins are

filled by an upwards shallowing Late Miocene to

Pleistocene sequence, commencing with deep¬

water pelagic sediments which are followed by

alluvial-fluvial and finally lacustrine deposits. In

the Rharb Basin, these sediments were deposited in

Paris ISBN: 2-85653-507-0. ,

This article includes 4 enclosures on 2 folded sheets.

88

M. ZIZI: ALPINE INVERSION, NORTH MOROCCO

Ceuta

Tanger

Mellilia

Guercita

Rabat

Tamlel

Boudenib

rrachidia

Internal zone

Flysch nappes

Mesorif

Intrarif

Miocene -pliocene sediments

(foredeeps and satellite basins)

unfolded atlantic

Meso-Cenozoic cover

Prerifain nappe

Folded Mesozoic

Atlas system and Rides Prenfames)

Hercynian granites " x Ma

Triassic

Nappe fronts

z/1 Paleozoic basement

Subsurface leading edge

(Prenfane nappe)

Rides Prerifaines

Guercif basin

FIG. I. Major tectonic element of Northern Morocco, showing location of study

areas (modified after Michard, 1975)

Source : MNHN , Paris

PER1-TETHYS MEMOIR 2: ALPINE BASINS AND FORELANDS

89

extensional satellite basins on top of the chaotic

“Nappe Prerifaine” (Flinch, 1993, this volume).

The “Rides Prerifaines”, translated from the

French as the “Fore-Rif ridges", form the moun¬

tainous terrain which extends from the Sidi Kacem

to Fes. These mountains are upheld by folded and

thrust-faulted shallow water Mesozoic carbonates

which are capped by Cenozoic sediments and the

chaotic Prerifaine nappe. Most authors interpreted

the “Rides Prerifaines” as south-west, south and

east verging thrust sheets (Sutter, 1980a and

1980b).

The southwest-northeast striking Middle Atlas

Mountains are located to the southeast of the Neo¬

gene Rif foreland. These mountains developed in

response to Late Cretaceous and Paleogene inver¬

sion of a system of deep Triassic to Jurassic half-

grabens. To the southeast, the Middle Atlas is

bordered by the High Plateau, a stable block which

is characterized by a relatively thin Mesozoic sedi¬

mentary cover, consisting of Triassic-Early Lias

redbeds and volcanics and Jurassic carbonates,

which rests on peneplained Palaeozoic rocks. The

Guercif Basin is a Late Neogene depression which

is superimposed on the northeastern part ot the

inverted Middle Atlas Trough.

STRATIGRAPHIC FRAMEWORK

The Mesozoic and Cenozoic sediments of

northwestern Morocco rest unconformably on the

peneplained surface of the Hercynian basement. As

evident from outcrops on the Moroccan Meseta,

this basement consists of deformed Carboniferous,

Devonian and Cambrian elastics and carbonates

which are intruded by Hercynian granites (Pique,

1982; Laville and Pique, 1991). During latest Car¬

boniferous and Permian times, accumulation of

continental elastics in small intramontane basins

was accompanied by the intrusion of acidic plutons

(Cousminer and Manspeizer, 1977; Van Houten,

1977).

During the Triassic and Early Jurassic, the

evolution of northern Morocco was dominated by

rifting activity which was intimately related to the

early phases of the Pangea breakup (Ziegler, 1988;

Dercourt et al., 1993). Extensional tectonics,

accompanied by the intrusion and extrusion of

tholeiitic basalts, controlled the subsidence of the

Middle and High Atlas Troughs and the external

Rif system of grabens and half-grabens

(Beauchamp, 1988; Favre et al., 1991; Laville and

Pique, 1991). Marine transgression entered these

grabens during the Late Triassic, giving rise to the

accumulation of a thick evaporitic series which lat¬

erally grades into continental red beds (Fig. 2).

During the Early Jurassic, open marine conditions

were established; whereas deeper water shales and

carbonates accumulated in the continuously sub¬

siding grabens, carbonate platforms developed on

the graben flanks (Favre and Stampfli, 1992). Tri¬

assic and Early Jurassic sediments range in thick¬

ness between 200 m and 2000 m.

With the early Middle Jurassic onset of sea

floor spreading in the Central Atlantic (Emery and

Uchupi, 1984), rifting activity ceased in Morocco.

However, continued tectonic activity, resulting in

the subsidence of small transtensional basins, must

be related to the sinistral translation of Africa rela¬

tive to Europe in response to progressive opening

of the Central Atlantic and the Western Tethys

(Ziegler, 1988; Laville and Pique, 1991; Dercourt

et al., 1993). During the Middle and Late Jurassic

times, the grabens of Morocco were gradually

filled in with elastics derived from southern

sources and later by platform carbonates, as evi¬

dent by the stratigraphic record of the Middle

Atlas, Guercif and the external Rif basins (Fig. 2).

With the Late Senonian onset of counter¬

clockwise convergence of Africa-Arabia with

Eurasia, the Alboran-Kabylia Block began to move

westwards with respect to North Africa and started

to converge with Iberia and northwestern Africa

(Wildi, 1983; Ziegler. 1988, 1990). Paleocene-

Early Eocene collision of the Alboran Block with

the Moroccan Tethys margin was accompanied by

the development of an accretionary wedge, corre¬

sponding to the Rif flysch nappes, the subsidence

of the Rif foreland basin and inversion of the Mid¬

dle Atlas Trough.

In the distal parts of the Rif foreland basin,

corresponding to the domain of the “Rides Preri-

faines", Middle Miocene (Langian-Serravalian)

shallow water carbonate and clastic sediments

transgressed over truncated Mesozoic strata; in

90

M. ZIZ1: ALPINE INVERSION, NORTH MOROCCO

RIDES

EAST

M ATLAS

PRERIFAINES

REKKAME

MASGOUT

TAOURIRT

SE AREA

PLIOCENE

MESSINAIN

TORTONIAN

LAN- SERRAVALIAN

PALEOCENE

TURONIAN

CENOMANIAN

LOWER

CRETACEOUS

PORTLANDIAN

KIMMERIDGIAN

CALLOVO-OXFOR

BATHONIAN

DOMERIAN

LOTHARINGIAN

BAJOCIAN

AALENIAN

TOARCIAN

TRIASSIC

PLIENSBACHIAN

FIG. 2. Chronostratigraphy of post-Hcrcynian series of Norlhern Morocco ("Rides

Prerifaincs,\ Middle Atlas and Guercif Basin).

Source : MNHN , Paris

PERI-TETHYS MEMOIR 2: ALPINE BASINS AND FORELANDS

91

turn, these are covered by Tortonian basinal marls.

During their accumulation, the Prerifaine nappe

was emplaced. This nappe consists of an up 2000

m thick chaotic assemblage of blocks, varying in

size from 1 m to 100 m, embedded in Tortonian

marls. Most common components are Triassic red

beds, evaporites and volcanics. In turn, the Preri¬

faine nappe is covered by Tortonian-Messinian

blue marls which grade upwards into the Early

Pliocene epicontinental sands and Late Pliocene

lacustrine limestones. Accumulation of this post¬

nappe sequence, which attains thicknesses of up to

1500 m, was contemporaneous with Messinian-

Pliocene foreland compressional phases. Quater¬

nary travertines, conglomerates, and yellowish

sands of presumably Villafranchian age, rest dis-

conformably on truncated Pliocene strata.

During these late deformation phases, Triassic

salt provided a regional detachment level. The

Early Mesozoic rift geometry and the distribution

of the Triassic salt played an important role in

guiding the geometry of the developing thrust

structures.

In the Guercif Basin, Tortonian to Pliocene

alluvial, deltaic and coastal sediments, reaching a

thickness of up to 2000 m, were deposited under a

tensional regime during Tortonian times and under

a compressional regime in Pliocene times. This

basin is superimposed on the northeastern parts of

the inverted Middle Atlas Trough.

This paper integrates the surface geology of

the “Rides Prerifaines” and the Guercif Basin with

reflection-seismic data imaging the subsurface

structures and documenting the larger scale geome¬

try and the areal extend of these features. Our

interpretation are based on seismic grids which

cover the entire “Rides Prerifaines and the Guer¬

cif Basin. Selected lines of these grids and their

interpretation are provided by Enclosures 1 to 4.

STRUCTURE AND EVOLUTION THE

“RIDES PRERIFAINES"

Surface geological studies ol the “Rides Preri¬

faines” date back to the late 1920’s. Daguin (1927),

Levy and Tilloy (1952). Durand Delga et al. (1960,

1962) and Sutter (1980a and 1980b) all suggested a

compressional origin for the “Rides Prerifaines”.

On the other hand, Faugere (1978) proposed that

these structures resulted from the interactions of

two basement-involving strike-slip systems, locat¬

ed along the southern and western margin of the

“Rides Prerifaines”. According to this author,

movements along these faults were transmitted to

the Mesozoic series which, due to their decoupling

from the basement by the Triassic salt, display

more complex structures. All authors agree that the

formation of the “Rides Prerifaines” is of Messin-

ian to Plio-Pleistocene age and post-dates the

major tectonic phases which controlled the evolu¬

tion of the Rif fold and thrust-belt.

Ait Brahim and Chotin (1984) carried out a

microtectonic study of the “Rides Prerifaines” and

identified four compressional phases, characterized

by principal horizontal compressional stress trajec¬

tories changing from NW-SE during pre- Miocene

times, to N-S during the Late Tortonian, to E-W

during the Messinian and to NE-SW during the

Plio-Pleistocene.

Figure 3 provides a tectonic map of the “Rides

Prerifaines”, showing their main structural ele¬

ments and the location of reflection-seismic lines

discussed below. The autochthonous foreland of

the Saiss plain lies to the south of the “Rides Preri¬

faines”. The northeastern parts of the “Rides Preri¬

faines” are overridden by the chaotic Prerifaine

nappe on which the Late Miocene to Pleistocene

post-nappe satellite basins subsided (Flinch, 1993,

this volume).

The seismic profiles, given in Enclosures 1

and 2, show that the dominantly NE-SW and E-W

trending surface structures are superimposed on

Early Mesozoic extensional fault systems. Seismic

line P-12 (Enclosure 1) crosses the “Rides Preri¬

faines” in an east-westerly direction and covers the

eastern part of the Rharb Basin, the Bou Draa,

Tselfat and Mesrana anticlines and the northern

parts of the Nzala des Oudayas structure. This pro¬

file shows that the “Rides Prerifaines” consist of

two large, superimposed thrust sheets, namely the

Prerifaine nappe and the Late Miocene-Pliocene

thrust belt w'hich involves the sedimentary fill of

the Mesozoic grabens, the “Aquitanian-Burdi-

galian” (Languian-Serravalian, according to recent

palaeontologic studies) foreland basin, as well as

92

M. ZIZI: ALPINE INVERSION, NORTH MOROCCO

LEGEND

f~1 Prerifain nappe (Accretionary wedge)

n Aquitanian-Burdigalian

f~n Oligocene

[23 Cretaceous

I I Dogger 4 _ >• thrusts faults

m upper Lias

EZ1 Lower Lias ^ ^ subsurface leading edge

■ Triasssic of the accretionary wedge

R H A R B

BASIN

SIDI KACEM

V

S A I S S

BASIN

MEKNES

10 kr

FIG. 3. Tectonic map of "Rides Prerifaines", showing location of seismic lines

given in Enclosures 1 and 2 and Fig. 5.

PERI-TETHYS MEMOIR 2: ALPINE BASINS AND FORELANDS

93

the Prerifaine nappe. This late compressional

deformation was contemporaneous with the depo¬

sition of the Late Miocene-Pliocene sedimentary

series (line P-15, Enclosure 2). The entire sedimen¬

tary package, including the allochthonous accre¬

tionary wedge of the Prerifaine nappe, is detached

from the autochthonous basement at the level of

Triassic evaporites.

The thrusted Bou Draa and Tselfat ridges are

superimposed on Triassic normal faults, offsetting

the top of the basement. The broad and the gentle

Mesarna structure, defined at intra-Jurassic levels,

is associated with lateral thicknesses changes

which are more pronounced at Toarcian and

Aaleno-Bajocian levels than in the Domerian:

these thickness changes are interpreted as reflect¬

ing intra-Jurassic salt movements. However, as the

base of the Prerifaine nappe is also deformed, a

Neogene growth component can be postulated for

the Mesarna structure. To the South of line P-12,

where Triassic salts are involved in this structure,

unconformities within the Late Mio-Pliocene sedi¬

ments give good evidences for reactivation of salt

movements at the time of the development of the

thrusted Bou Draa and Tselfat structures.

The northwest-verging the Bou Draa structure

is evidently associated with a sharp increase in

thickness of the Triassic-Jurassic sediments across

a deep seated basement fault. This fault is part of

the Sidi Fili fault system which, decades ago. had

been identified by petroleum geologists.

The seismic profile P-15 (Enclosure 2), which

94

M. ZIZI: ALPINE INVERSION, NORTH MOROCCO

WNW

ESE

s

JEBEL OUTITA

WESTERN FLANK

4 Km

ABREVIATIONS:

TR : TRIASSIC ; P-DO : PREDOMERIAN ; DO : DOMERIAN ; TO : TOARCIAN

Aa-Baj : AALENO- BAJOCIAN ; LAN-SERR : LANGHIAN- SERRAVALIAN

NP : NAPPE PRERIFAINE ; UM-P : UPPER MIOCENE-PLIOCENE

FIG. 5. E-\\ profile throgh "Rides Prerifaines". illustrating block-faulted basement, sub-

honzonial reflectors beneath intra-salt decollement horizon, progradation to the cast of the

Aaleno-Bajocian sequence, convergence of Late Miocene-Pliocene sequence towards the

western flank of the Bou Kennfoud structure.

Source : MNHN . Paris

SIC)

PER1-TETHYS MEMOIR 2: ALPINE BASINS AND FORELANDS

95

Source : MNHN , Paris

FIG. 6. Geological map of Guercif Basin, showing location of outcropping diapirs and seismic line given in Fig. 7.

96

M. ZIZI: ALPINE INVERSION. NORTH MOROCCO

crosses the “Rides Prerifaines” in a northeasterly

direction, shows that also the southerly verging

thrust faults, which carry the Kefs and Kheloua

structures, ramp up from the Triassic evaporites

through Jurassic strata above a set of normal base¬

ment faults. Also the largely salt induced Nzala des

Oudaya structures are associated with Triassic nor¬

mal faults offsetting the basement (Enclosure 1).

The Triassic-Jurassic fault map, given in

Fig. 4, is based on an interpretation of the entire

seismic grid. It illustrate that the Early Mesozoic

fault system consists of west, east and north hading

normal faults, the length of which varies between

2 km and 20 km. The Sidi Fili fault trend defines

northwestern margin of the Triassic-Early Jurassic

“Rides Prerifaines” half graben system.

On seismic lines, the Prerifaine nappe is char¬

acterized by discontinuous to chaotic reflectivity,

abounding with diffractions (Fig. 5). Its southern

termination is clearly evident on the line P-15

(Enclosure 2); this allochthonous body is laterally

offset and onlapped by Tortonian series, character¬

ized by parallel and laterally continuous reflectors.

The base of the nappe corresponds to the smooth

surface which forms the top of the “Aquitanian-

Burdigalian” reflectors. This clearly shows that

this allochthonous unit was emplaced in post-Mid-

dle Miocene and pre-Late Miocene times (Flinch,

1993, this volume), that is, during a relatively short

time span. Emplacement of this nappe was accom¬

panied by drowning out of the foreland platform

and its rapid subsidence to considerable water

depth.

The Late Miocene to Pliocene series, which

covers the Prerifaine nappe, is characterized paral¬

lel to sub-parallel reflectors which generally con¬

verge towards the top of “Rides Prerifaines”

anticlines (Enclosure 2). Significant syn-deposi-

tional deformation is indicated by the convergence

of reflectors and the presence of unconformities

within this Mio-Pliocene series, as shown by the

growth of the syncline away from the anticlines

(Enclosure 2). Moreover, the thrust fault carrying

the Bou Draa structure (Enclosure 1) clearly cuts

through the Prerifaine nappe and the Late Mio-

Pliocene series. Therefore, deformation of the

“Rides Prerifaines” clearly post-dates the emplace¬

ment of the Prerifaine nappe. This was already evi¬

dent from surface geology (e.g. Levy and Tilloy,

1952; Sutter, 1980).

The northerly striking structures of “Rides

Prerifaines” are generally associated with pre¬

existing salt pillows, some of which are superim¬

posed on normal basement faults. Line P-12 and

P-15 show that during the Late Miocene-Pliocene

deformation of the “Rides Prerifaines” the Meso¬

zoic and younger series were decoupled from the

basement at the level of Triassic salts. Although

there is no evidence for the reactivation of normal

faults affecting the basement, their intra sedimenta¬

ry part was clearly reactivated and played an

important role in localizing the deformation of

structures which now form the “Rides Prerifaines”

(Fig. 5). As such, the entire system of the “Rides

Prerifaines” must be considered as a thin-skinned

thrust belt which partly scooped out the sedimenta¬

ry fill of the Triassic-Early Jurassic Prerifaine

grabens. The thrusted folds forming the western

and eastern structures correspond to lateral ramps

whereas the southern structures form the frontal

ramps of this thin-skinned thrust belt.

EVOLUTION OF THE GUERCIF BASIN

The Guercif Basin contains up to 2000 m of

Tortonian to Pliocene sediments and is superim¬

posed on the northeastern parts of the Early Meso¬

zoic Middle Atlas Trough which was inverted

during Paleogene times (Figs. 6 and 7).

Development of the Middle Atlas Mountains

is generally thought to result from the sinistral

transpressional deformation from an Early Meso¬

zoic rifted basin (Mattauer et al., 1977; Jacobsha-

gen, 1988; Fedan, 1988; Boccaletti et al., 1990;

Bernini et al., 1994). Choubert and Faure-Muret

(1962) visualize Late Eocene, Late Oligocene and

Middle Miocene compressional phases whereas du

Dresnay (1988) recognized Late Senonian precur¬

sor events followed by a major Late Eocene phase

of basin inversion.

The tectonic history of the Neogene Guercif

Basin was studied by Colletta (1977) who identi¬

fied a Late Tortonian to Messinian extensional

phase, a late Pliocene compressional episode and

possibly renewed extension during the Quaternary.

Source : MNHN . Paris

GRF1

PERI-TETHYS MEMOIR 2: ALPINE BASINS AND FORELANDS

97

LJJ

(/)

LU

5

z

5

o

c

‘■J

o

_o

K,

“O

c

r3

c

O

.1^

n a

O v)

n c

o o

H -

C3

SO O

c O

= 2

|S

if

Ml

o o

— c

2 2

-c SO

C ■-

• — -

c

M n

I H

*C 'c

•s.

g §

• — V)

I §

i!

^ E

2 s

u. Z

Source : MNHN. Pans

98

M. ZIZI: ALPINE INVERSION. NORTH MOROCCO

Mokhtari (1990), Boccaletti et al. (1990) and

Bernini et al. (1994) all interpreted the Guercif

Basin structures in terms of flower structures.

Bally (1992) described the structures of the Guer¬

cif Basin as mini-inversion structures, involving

the reactivation of Mesozoic extensional faults dur¬

ing pre-Tortonian inversion movements, however,

without fully compensating their Mesozoic offsets.

The following discussion on the evolution of

the Guercif area is based on the analysis of a grid

of reflection-seismic profiles which are partly cali¬

brated by wells.

Lack of well control impedes precise dating of

the presumably Late Triassic-earliest Jurassic onset

of the rifting activity in the northeastern parts of

the Middle Atlas Trough ( line G-17, Enclosure 3).

Strong thickness changes across the fault limiting

the Debdou platforms suggests that differential

subsidence of the Middle Atlas half-graben persist¬

ed into early Late Jurassic times; diverging intra-

Jurassic reflectors and a number of unconformities

are taken as evidence for syn-sedimentary exten¬

sion. An unconformity at the base of the Bathonian

to Late Jurassic sequence indicates continued tec¬

tonic instability of the area.

Some salt-cored anticlines have been mapped

in the western Guercif Basin (Fig. 6). Line G-5

(Enclosure 4) gives a good impression of these

structures even though the seismic data are not

very good. This section crosses the outcropping

Rhorgia diapir to the SW and the Bou Msaad diapir

to the East. The presence of a third diapir between

these two structures is suggested. Note the inter¬

preted thinning of the lower Liassic-Domerian

interval towards the western domes and the thick¬

ening of the Bathonian-Portlandian section along

the southwestern flank of the central salt structure.

The thinning of the Early Jurassic intervals can be

interpreted as reflecting the pillow phase of the

diapirs which was followed by salt evacuation dur¬

ing the Bathonian-Portlandian, the rise of the diapir

and the formation of a rim syncline during the Late

Jurassic. Therefore, it is concluded that halokinetic

movements were initiated during the Early Juras¬

sic, resulting in the formation of salt pillows, and

culminated by the end of the Jurassic with the rise

of the diapirs and the development of rim syn¬

clines.

In the area of the Guercif Basin, a major hia¬

tus spans Cretaceous to Middle Miocene times.

However, in the southeastern part of the Middle

Atlas, Cenomanian to Maastrichtian sediments are

preserved (Fig. 2). Enclosure 3 clearly illustrated

that the Mesozoic half graben, which underlies the

Neogene Guercif Basin, was only mildly inverted

prior to the transgression of the Tortonian and

younger series. During these pre-Neogene inver¬

sion movements, the basement fault bounding the

Debdou platform was reactivated; transpressional

movements along this fault gave rise to the devel¬

opment of an anticlinal structure in the hanging-

wall block, causing uplift and erosion of its

Bathonian to Portlandian sedimentary cover.

Surface geological data from the southeastern

Middle Atlas indicates the occurrence of pre-

Maastrichtian inversion events (du Dresnay, 1988),

which were followed by the commonly accepted

Late Eocene event (Michard, 1976; Robillard,

1979; Ziegler, 1988). Therefore, it is reasonable to

assume that the main inversion of the northeastern

part of the Middle Atlas Trough had also occurred

during the Late Senonian to Late Eocene time

span. However, the extent to which this area had

been covered by Late Cretaceous sediments prior

to its inversions is unknown.

Subsidence of the Guercif Basin commenced

with the transgression of Tortonian strata over trun¬

cated Jurassic series and persisted, under regres¬

sive conditions, through Messinian into Pliocene

times. Basal conglomerates are followed by lacus¬

trine shale and carbonates which are capped by

Pliocene continental sands (Fig. 2).

The seismic profiles across the Guercif Basin,

given in Fig. 7, show that its Tortonian subsidence

was governed by extensional tectonics. As Messin¬

ian and Pliocene strata are not affected by tensional

faults, rifting activity must have been of relatively

short duration. The reflection configuration of

Messinian strata indicates that they were deposited

during a tectonically quiescent period. Conver¬

gence of Pliocene reflectors over structures, such

as the one drilled by the well GRF1 (Fig. 7) and

the Safsafat anticline, indicate that a compressional

regime dominated the Pliocene evolution of the

basin. During the Pliocene partial inversion of the

basin, some of the Tortonian extensional faults

were apparently compressionally reactivated. Dur¬

ing this late compressional phase, the diapirs were

reactivated to the extent that they now form out¬

cropping elongated folds.

Source : MNHN . Paris

PERI-TETHYS MEMOIR 2: ALPINE BASINS AND FORELANDS

99

It is concluded that the area occupied by the

Neogene Guercif Basin has undergone a plyphase

evolution. Triassic to Early Jurassic crustal exten¬

sion governed the subsidence of the Middle Atlas

Trough. Tectonic instability continued during Mid¬

dle and Late Jurassic times. Transpressional defor¬

mation of the Middle Atlas Trough commenced

during the Late Senonian and culminated in its

Eocene inversion. Tortonian extension governed

the subsidence of the Guercif Basin. Messinian

strata were deposited under a quiescent regime.

Remobilisation of the Triassic salts under the over¬

burden of thick Neogene sediments gave rise to the

development of diapirs and domal structures.

These features were modified during the Pliocene

compressional phase. Pliocene compressional

structures strike NE-SW and N-S.

CONCLUSIONS AND IMPLICATIONS FOR

HYDROCARBON EXPLORATION

The Triassic-Jurassic grabens of northern

Morocco form an integral part of the Western

Tethys and North and Central Atlantic rift system.

During the Alpine orogeny these grabens were

inverted to various degrees. Eocene inversion of

the Middle Atlas rift must be related to collisional

coupling between the evolving Rif orogen and its

foreland. The emplacement of the Prerifaine nappe

at the transition from the Middle to the Late

Miocene was accompanied by rapid subsidence of

a relatively narrow foreland basin. Late orogenic

phases of foreland compression resulted in the

destruction of this foreland basin. Eocene inversion

structures of the Guercif Basin involve compres¬

sional reactivation of tensional basement faults. In

contrast, there is no evidence for basement reacti¬

vation during the development of the Messinian-

Pliocene “Rides Prerifaines”; these are

characterized by a major detachment system which

soles out in Triassic evaporites, deposited in an

extensional basin. South-verging thrusts are associ¬

ated with east- and west-verging lateral ramps. All

structures of the “Rides Prerifaines" are strongly

influenced by the configuration of the friassic-

Early Jurassic rifted basin, the sedimentary fill of

which was partly scooped out by Messinian-

Pliocene thrust faulting.

The “Rides Prerifaines” contain significant

hydrocarbon accumulations, contained in Late

Neogene structures. During the Messinian-

Pliocene deformation of the “Rides Prerifaines”,

hydrocarbons contained in pre-existing salt

induced structural traps, were partly destroyed,

thus releasing sizable volumes of oil for the charge

of the newly formed structural traps. Early salt

induced structures, which retained part of their clo¬

sure during the late phases of basin deformation,

may still contain significant amounts of hydrocar¬

bons.

In the Guercif Basin, four exploratory wells

were drilled on Pliocene structures all failed to

encounter hydrocarbons. Eocene inversion struc¬

tures are more likely to contain hydrocarbon in

Mesozoic reservoirs, provided Early Jurassic

source-rocks have attained maturity, as attested by

oil seeps found the Middle Atlas (e.g. Issouka oil

seep). Triassic-Jurassic extensional tectonics con¬

trolled the distribution of reservoirs, seals, source-

rocks in half grabens and migration pathways of

hydrocarbon generated. Generally, the footwall is

dominated by relatively shallow water facies, such

as carbonate build-ups, sand shoals and slope tur-

bidites. The hanging wall is marked by ramp-type

margin facies, including sand shoals and reefs.

Palaeotectonic basin analyses, based on an

integration of surface and subsurface geological

data, as well as structural and seismostratigraphic

analyses of reflection seismic data, can greatly

advance the understanding of the evolution of fore¬

land basins and their hydrocarbon habitat.

Acknowledgments^ The interpretations pre¬

sented here were developed during the preparation

of the author’s Ph.D. thesis at Rice University,

Houston , under the supervision of Prof. A.W.Bally,

to whom he wishes to expresses his sincere thanks.

The support of this research project by ONAREP ,

which provided the seismic and well data, and by

TOTAL , which financed it, is gratefully acknowl¬

edged. Thanks are extended to Dr. P.A. Ziegler for

his constructive and critical review of an earlier

version of this manuscript and his editorial efforts.

100

M. ZIZI: ALPINE INVERSION. NORTH MOROCCO

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Enclosures

Enclosure 1 Seismic line P-12. Rides Prerifaines

Enclosure 2 Seismic line P-15, Rides Prerifaines

Enclosure 3 Seismic line G- 1 7, Guercif Basin

Enclosure 4 Seismic line G-5, Guercif Basin

Source : MNHN, Paris

The Valencia Trough:

geological and geophysical constraints

on basin formation models

M. Torne, E. Banda & M. Fernandez

Institute of Earth Sciences (J. Almera).

Consejo Superior de Investigaciones Cientificas,

Lluis Sole i Sabaris s/n, E-08028-BarceIona, Spain

ABSTRACT

The Valencia Trough, located between the

Spanish mainland and the Balearic Promontory, is

one of the Cenozoic extensional basins of the

Western Mediterranean which developed in a

region of convergence between the European and

African plates. The Valencia Trough began to sub¬

side during the late Oligocene-early Miocene along

the Mediterranean side of Spain, coeval with a

phase of compression which affected its southeast¬

ern margin, formed by the Balearic Promontory.

This rifting phase was followed by a period of

post-rift subsidence. During the Plio-Quaternary,

extensional tectonics and minor volcanic activity

resumed along the Spanish margin of the Trough.

This phase of extension is, however, not recorded

in the central parts of the basin. The wealth of geo¬

logical and geophysical data collected from the

Valencia Trough during the last decades provides a

precise image of its crustal structure; however, its

deep lithospheric configuration is still poorly

known.

The Valencia Trough is characterized by a

strongly attenuated continental crust, except in its

northeasternmost part where oceanic crust is prob¬

ably present. It is underlain by an anomalous low-

velocity uppermost mantle, the lateral extent and

thickness of which are still under debate. Kinemat¬

ic models of basin development have been quite

successful in describing some aspects of the central

parts of the Trough, though not of the entire basin

and particularly not for its southern region. As

these models failed to integrate the complex plate

interactions which governed the development of

the Valencia Trough, they must be regarded as sim¬

plistic. The Valencia Trough hosts an oil province

which is largely restricted to the Ebro delta.

INTRODUCTION

The Valencia Trough is a NE-SW oriented tri¬

angular shaped basin which is located between the

Spanish mainland and the northeastern prolonga¬

tion of the Betic Cordillera, the Balearic Promonto¬

ry (Fig. 1). The continental crust underlying this

trough was consolidated during the Hercynian

orogeny. During the Mesozoic break-up of Pangea,

Torne M Banda E & Fernandez. M., 1996. — The Valencia Trough: geological and geophysical constraints on basin formation

models. In: Ziegi’er. P. A.’& Horvath. F. (eds). Peri-Tethys Memoir 2: Structure and Prospects of Alpine Basins and Forelands. Mem. Mus.

nain. Hist. not.. 170: 103-128. Paris ISBN : 2-85653-507-0.

104

M. TORNE. E. BANDA & M. FERNANDEZ: VALENCIA TROUGH. SPAIN

Alpine Thrust Belts [ . : ! • | Cenozoic Rift Basins

Cenozoic Oceanic Crust

FIG. 1. Location map of the study area showing main geological features of the

Western Mediterranean (modified after Banda and Santanach. 1992b). Valencia

Trough is outlined by heavy grey line. C.C.R.: Catalan Coastal Ranges. Black

squares: calcalkaline magmatism (30-15 Ma); Open circles: alkaline magmatism

(15-0 Ma). (location of volcanic outcrops from Marti et al.,1992).

this crust underwent repeated extensional events,

controlling the subsidence of graben structures.

During the latest Cretaceous and Paleogene,

intraplate compressional stresses caused inversion

of the Iberian Chain, the Catalan Coastal Ranges

and their off-shore equivalents (Fig. 1); during this

process, Mesozoic extensional crustal thinning was

apparently largely recovered.

The present day structure of the Valencia

Trough is thought to result from a late Oligocene-

early Miocene rifting event which was coeval with

a phase of compression confined to the southeast¬

ern margin of the basin, formed by the Balearic

Promontory. This extensional phase affected main¬

ly the northwestern part of the basin, as evident by

the development of a series of horst and graben

structures, bounded by ENE-WSW trending nor¬

mal faults. Whereas the horsts are held up by a

variety of Palaeozoic and Mesozoic rocks, the

grabens were filled by late Oligocene-early

Miocene shales (Torres and Bois, 1993). In con¬

trast, the Balearic part of the basin is characterized

by a series of SE-NW trending thrusts and reverse

faults which developed during a Late Oligocene to

Middle Miocene compressional phase (Roca and

Desegaulx, 1992). Subsequently, these compres¬

sional structures were tensionally reactivated,

resulting in the development of horst and graben

structures, as seen on the Island of Mallorca

(Fig. 1). For further details the reader is referred to

Banda and Santanach (1992a).

Source : MNHN. Paris

PERI-TETHYS MEMOIR 2: ALPINE BASINS AND FORELANDS

105

The Valencia Trough has been the target of

extensive geological and geophysical investiga¬

tions by the petroleum industry and academic insti¬

tutions (Fig. 2). During the 1970's and 1980's, the

Catalonian shelf was intensely explored for hydro¬

carbons. The recording of extensive reflection-seis¬

mic surveys and the drilling of over 90 wells

(Fig. 2a) resulted in the discovery of a number of

small and one large oil fields, having cumulative

ultimate recoverable reserved of the order of 250 to

300 x 106 bbls of oil. Most of these accumulations

are contained in extensional fault blocks and

buried hills, upheld by karstified Mesozoic carbon¬

ates. These structures are sealed by middle

Miocene basinal shales. Hydrocarbon charge is

provided by Mesozoic source-rocks as well as by

late Oligocene-early Miocene shales which were

deposited during the early rifting stage of the

Valencia Trough. These data sets provide valuable

information on the sedimentary record of the

Valencia Trough and its subsidence and thermal

evolution.

The first academic geophysical experiments

commenced in the early 70's with the acquisition

of two seismic refraction profiles, located North

and South of the Island of Mallorca (Hinz, 1972;

Gobert et al., 1972), followed up by a seismic

refraction/wide-angle reflection experiment along

the Balearic Promontory (Banda et al., 1980). Dur¬

ing 1988 the VALSIS experiments were carried out

to determine the lithospheric configuration of the

basin. During a first cruise (VALSIS-I) heat-flow

measurements were acquired along four transects

crossing the axis of the Trough (Fig. 2b; Foucher et

al., 1992). During a second cruise (VALSIS-II),

twelve deep multichannel seismic-reflection pro¬

files (CDP), eight wide-aperture multichannel pro¬

files (COP) and six expanded spread profiles (ESP)

were recorded (Figs. 2c and 2d). Marine seismic

data were complemented by land recording of

shots (Gallart et al., 1990). An additional refrac¬

tion/wide-angle reflection experiment was carried

out during the summer of 1989 along three profiles

crossing the basin (P-I, P-II, and P-III of Fig. 2d;

Danobeitia et al.. 1992). For a detailed discussion

of CDP results data the reader is referred to Mauf-

fret et al. (1992), Maillard et al. (1992) and Torne

et al. (1992). The ESP results are discussed by Pas¬

cal et al. (1992) and Torne et al. (1992) and the

wide-angle data by Gallart et al. (1990) and

Danobeitia et al. (1992). COP results were present¬

ed by Collier et al. (1994). In 1992, the crustal

structure of the Valencia Trough was investigated

by coincident steep and wide-angle reflection data

(Fig. 2c) under the auspices of the Spanish Estu-

dios Sismicos de la Corteza Iberica (ESCI) Pro¬

gram (Gallart et al. 1995; Vidal et al., 1995).

The results of these experiments show that the

Valencia Trough is characterized by a strongly

attenuated continental crust which is underlain by

an anomalous low-velocity upper mantle (7.6 to

8.0 km/s), except in its northeasternmost part

where oceanic crust is probably present. In the cen¬

tral part of the basin, the crust has a thickness of

about 15-16 km; towards its margins it increases

asymmetrically to average values of 20-22 km

along the Iberian coast and about 24 km below

Mallorca. Along the flanks of the basin, the upper

and middle crust are characterized by an almost

transparent layer with velocities ranging from 6.1

to 6.4 km/s. In contrast, the lower crust is variably

reflective below the flanks of the basin and has

velocities in the 6.4 to 6.9 km/s range, whereas it is

almost absent under its axial parts.

Information on the configuration of the lithos-

phere-asthenosphere boundary comes primarily

from modelling results. Modelling of the central

part of the basin shows that the lithosphere thins

towards the axis of the basin to values in the 60 to

65 km range (e.g.. Watts and Tome 1992a; Zeyen

and Fernandez, 1994). Moreover, surface- wave

studies favour thinning of the lithosphere towards

the central part of the basin (Marillier and Mueller,

1985). There is no information, however, on the

position of the lithosphere-asthenosphere boundary

in the southern parts of the Trough, and also in the

North at its transition to the oceanic Provencal

basin (Fig. 1).

Direct evidence for volcanic activity comes

from outcrops and exploration wells (Lanaja,

1987) and DSDP Site 123 (Ryan et al., 1972).

Reflection-seismic and aeromagnetic data provide

indirect evidence for additional volcanic centres

(e.g., Mauffret, 1976; Maillard et al., 1992; Marti

et al., 1992; Galdeano et al., 1974) (Fig. 1). Fol¬

lowing Marti et al. (1992), Cenozoic magmatism in

the area is characterized by two volcanic cycles

which are clearly separated in time and by their

petrology and tectonic setting. The first volcanic

cycle (late Chattian-early Burdigalian) coincides

106

M. TORNE, E. BANDA & M. FERNANDEZ: VALENCIA TROUGH, SPAIN

Precambrian and

Paleozoic rocks

Paleozoic, Mesozoic and Tertiary

rocks in Alpine Belts

Tertiary basins

FIG. 2. a) Location of hydrocarbon exploration wells along the Ebro platform

and Spanish margin. C.C.R.: Catalan Coastal Ranges; I.C.: Iberian Chain; E.I:

Island of Eivissa; Mn.I.: Island of Menorca; M.I.: Island of Mallorca.

b) Heal How determinations in mW/m^.

c) Track lines of deep multichannel seismic profiling. Thick lines: Valsis-II

CDP/COP profiles. Dashed-dotted line: ESCI coincident steep and wide-angle

reflection profile.

d) Thick lines: Valsis-II ESP Profiles (triangles show mid-point locations). Dashed-

dotted lines: wide-angle/refraction seismic profiles.

Source : MNHN , Paris

PERI-TETHYS MEMOIR 2: ALPINE BASINS AND FORELANDS

107

Precambrian and

Paleozoic rocks

Paleozoic, Mesozoic and Tertiary

rocks in Alpine Belts

Tertiary basins

FIG. 3. a) Bathymetry of Valencia Trough, 200 m contour intervals.

b) Free-Air gravity anomaly map based on all available marine gravity data. Con¬

tour interval 10 mGal. Abbreviations as in Fig. 2.

c) Simplified Bouguer gravity anomaly map, contour interval 10 mGal (after Torne

et al., 1992). Data on Iberia are based on Casas et al. (1987) and on Mallorca on

IGME (1981).

d) Geoid anomaly map, contour interval 0.5 m. The map is based on GEOMED

(e.g., Sevilla, 1992). A regional field based on the OSU9IA model to degree and

order 1 2 has been removed from the observed data prior gridding.

Source

108

M. TORNE, E. BANDA & M. FERNANDEZ: VALENCIA TROUGH. SPAIN

with the initial, main phase of rifting and is charac¬

terized by calcalkaline andesitic and pyroclastic

rocks. The second cycle (Tortonian to Recent) is

characterized by poorly differentiated alkali-basalts

and was accompanied by a minor phase of exten-

sional tectonics.

Paleomagnetic investigations suggest a 20°

early-middle Miocene clockwise rotations of the

islands of Mallorca and Menorca relative to the

Catalan Coastal Ranges, coinciding with the

emplacement of the Betic-Balearic thrust sheets. A

further late Miocene-Pliocene rotation of up to 20°

may be related to extensional faulting. However, it

can not be excluded that part of these rotations are

related to whole lithospheric movements (Pares et

al., 1992). These authors point out that clockwise

rotation in the Balearic Islands is in disagreement

with most of the models proposed for the evolution

of the Western Mediterranean, and that the total

amount of rotation is too large to be accounted for

by crustal extension in the Valencia Trough.

In summary, the various studies of the basin,

using different geological and geophysical

approaches, have resulted in a wide range of geo¬

dynamic hypotheses. These range from back-arc

extensional mechanisms, related to northwestward

subduction of oceanic lithosphere (e.g., Bocaletti

and Guazzone, 1974; Mauffret, 1976; Banda and

Channel, 1979; Cohen, 1980; Horvath and Berck-

hemer, 1982) to intra-continental rifting, related to

horizontal movements of crustal blocks and sea¬

floor spreading (e.g., Auzende et al., 1973; Rehault

et al., 1985) and rift propagation from the Alpine

forelands across the Alpine megasuture (Ziegler,

1988, 1992). Recently, Fontbote et al. (1989) and

Roca and Desegaulx (1992) postulated that devel¬

opment of the Valencia Trough could be explained

by foreland-type mechanisms, whereas Doblas and

Oyarzun (1990) proposed that its origin is con¬

trolled by a major extensional detachment surface,

cutting the crust and lithosphere, and upwelling of

the asthenosphere.

In this paper we summarize the present-day

crustal and lithospheric configuration of the Valen¬

cia Trough and focus on modelling results and dif¬

ficulties and limitations which have been

encountered when applying “classical” extensional

models to explain the origin and evolution of this

basin.

GEOPHYSICAL OBSERVATIONS

In this section we present a compilation of the

various geophysical data sets which provide con¬

straints on modelling the evolution of the Valencia

Trough (Figs. 2 and 3). Figure 3a illustrates the tri¬

angular shape of this basin and its opening to the

Northeast where water depths reach values of up to

2500 m, whereas in its southwestern parts they not

exceed 1200 m. The axial trough is flanked by nar¬

row shelves, except in the area of the Ebro delta

(Fig. 3a).

Up to the VALSIS-I survey, information on

the thermal regime of the area was limited to tem¬

perature gradients obtained from on-shore and off¬

shore water and oil wells. These indicate that areas

adjacent to the Trough are characterized by large

temperature gradient variations, probably related to

ground-water circulation (Fernandez and Banda,

1989; Fernandez et al., 1990). Similarly, heat-flow

values for the eastern part of the Ebro basin (Cabal

and Fernandez, 1995) and the Spanish shelf

(Negredo et al., 1995a) are also high variable and

range from 55 to 90 mW/m- The islands of Mal¬

lorca and Menorca are characterized by back¬

ground heat flow values of about 70-80 mW/m“

and 75-90 mW/m2, respectively (Fernandez and

Cabal, 1992). Results of the VALSIS-I survey

(Foucher et al., 1992), which are not corrected for

the thermal blanketing effect of sediments, indicate

for the axial part of the basin heat-flow values

varying from 88 mW/m2 in the Southwest to

66 mW/m2 in the Northeast at the transition to the

Provencal basin, and thus demonstrate a significant

decrease in heat-flow values towards deeper waters

(Fig- 2b).

Free-Air gravity data (Fig. 3b) show that the

axial parts of the Trough are dominated by values

around 0 mGal whereas on its flanks maximum

values of about 40 mGal are reached. This may

suggest that the regional features of the study area

are in local isostatic equilibrium; this is in accor¬

dance with various modelling results which show

that the lithosphere has acquired little or no

strength since rifting (e.g. Watts and Torne, 1992b;

Zeyen and Fernandez, 1994). Bouguer gravity data

shows that the Trough is associated with a gravity

anomaly high of about 100-150 mGal (Fig. 3c). 2D

Source : MNHN. Paris

PERI-TETHYS MEMOIR 2: ALPINE BASINS AND FORELANDS

109

gravity modelling by Watts and Tome (1992a)

shows that this gravity high can be explained by

the combined effect of mantle and crustal thinning.

Figure 3d gives a map of geoid anomalies

derived from GEOMED (e.g., Sevilla, 1992). Since

the geoid is more sensitive to deep mass distribu¬

tion than gravity, this can help in better decipher¬

ing the topography of the base of the lithosphere.

Figure 3d shows that the Valencia Trough is asso¬

ciated with a relatively broad negative geoid anom¬

aly which increases in magnitude toward the

northeastern, whereas both flanks, and particularly

the Balearic Promontory, are associated with posi¬

tive geoid anomalies of up to 2 m. A first evalua¬

tion of this geoid anomaly low favours the

hypothesis that the lithosphere is thinner under the

central part of the Trough.

CRUSTAL AND LITHOSPHERIC

STRUCTURE

In the following we discuss the sedimentary

fill of the Valencia Trough, its upper and lower

crustal configuration and the structure of the lithos-

phere/asthenosphere boundary.

Sedimentary Record and Basement Structure

Following Soler et al. (1983), the stratigraphic

record of the Valencia Trough and surrounding

areas is characterized by a major unconformity

separating the Palaeozoic-Mesozoic basement from

its Oligocene to Quaternary sedimentary cover

(Fig. 4). On the Spanish shelf, Hercynian deformed

Palaeozoic sedimentary and metamorphic rocks are

uncomformably overlain by Permian and Mesozoic

series, consisting of carbonate, siliciclastic and

evaporite rocks (Torres and Bois, 1993). Rapid lat¬

eral thickness variations and facies changes indi¬

cate that these sediments accumulated in tensional

basins which developed in conjunction with the

opening of the Tethys Ocean (e.g. Ziegler, 1988;

Fontbote et al., 1989; Maillard et ah, 1992). On the

islands of Mallorca and Eivissa (Fig. 2a), Middle

Jurassic-Late Cretaceous series were developed in

a slope and base of slope facies, reflecting their

location along the Tethys passive margin (Banda

and Santanach, 1992b).

The latest Cretaceous and Paleogene phase of

intraplate compression was responsible for the

inversion of the Mesozoic basins and uplift of the

present-day off-shore parts of the Valencia Trough

(Fig. 1). Inversion of the Iberian Chain was accom¬

panied with the development of a series of NW-SE

and E-W striking thrust faults (Guimera, 1984;

Guimera and Alvaro, 1990) whereas the Catalan

Coastal Ranges evolved in response to convergent

wrench movements along NE-SW striking faults

(Anadon et al., 1985). In the off-shore, Paleogene

rocks are generally absent, except in the Barcelona

graben (Bartrina et al., 1992) where continental or

transitional deposits are present. In contrast, wide¬

spread Eocene to Oligocene lacustrine carbonates

and lignites occur on the island of Mallorca

(Ramos-Guerrero et al„ 1989).

The Oligocene to Quaternary sedimentary fill

of the Valencia Trough can be subdivided in four

depositional sequences, separated by unconformi¬

ties (Fig. 4b; Garcfa-Sineriz et al.. 1979; Soler et

al., 1983; Anadon et al., 1989; Clavell and

Berastegui, 1991).

On the western flank of the Trough, the lower

sequence, spanning late Oligocene to Langhian

times, consists of basal continental-transitional

elastics and shales, which accumulated in fault-

bounded shallow basins, and the transpressive

early Miocene marine shales and carbonates of the

Alcanar formation. The latter was deposited after

the first rifting stage (Chattian-Aquitanian) and

during the Betics compressional phase (Torres and

Bois, 1993) and oversteps the block-faulted relief

of the Valencia rift (Fig. 5). On the southeastern

flank of the basin, Chattian-Aquitanian sediments

consist of shallow-water carbonate and clastic

rocks (Rodri'guez-Perea, 1984; Anglada and Serra-

Kiel, 1986) whereas Burdigalian and Langhian

series consist of pelagic and calcareous turbidites

that were deposited during the thrust deformation

of the Balearic Promontory.

The second sequence spans Serravallian to

early Messinian times. On the Balearic Promontory

it is represented by carbonates deposited under a

110

M. TORNE. E. BANDA & M. FERNANDEZ: VALENCIA TROUGH. SPAIN

ONSHORE

OFFSHORE

CATALAN-VALENCIAN SHELF

BALEARIC PROMONTORY

b)

CATALAN MARGIN

LEGEND

MALLORCA

FIG. 4. a) Simplified stratigraphic charts of Valencia Trough region

(after Banda and Santanach. 1992b). On-shore and off-shore columns

after Bartrina et al. (1992).

b) Stratigraphic columns of the Spanish margin and Mallorca (after Tor¬

res and Bois. 1993).

Source : MNHN . Paris

VALENCIA TROUGH

PERI-TETHYS MEMOIR 2: ALPINE BASINS AND FORELANDS

Hi

Source : MNHN. Paris

112

M. TORNE. E. BANDA & M. FERNANDEZ: VALENCIA TROUGH. SPAIN

tectonically quiet regime, whereas from the Span¬

ish coast, a major deltaic complex, referred as the

Castellon Group, prograded into the Valencia

Trough, characterized by rapid deepening in

response to post-rift thermal subsidence.

During the Messinian, a major drop in sea-

level gave rise to the development of high relief

unconformities on the continental shelves and the

deposition of variably thick halites and minor sul¬

phates in the central and northeastern parts of the

Trough (Mulder, 1973; Ziegler, 1988). Sequence

three is therefore only represented in the deepest

parts of the basin.

The forth sequence, which spans late Messin-

ian-Quaternary times, corresponds to the Ebro

Group which, similar to the Castellon Group, con¬

sists of a major deltaic complex building out into

deeper waters from the Spanish coast. However, it

records in coastal and shelf areas a renewed phase

of extensional tectonics that was accompanied by

the extrusion of alkaline volcanics, forming the

Columbretes Islands. Similar extensional faults are

no evident in the central parts of the Trough

(Banda and Santanach, 1992b).

Figure 6a provides a structural map at the base

of the Cenozoic sedimentary till ol the Valencia

Trough (Lanaja, 1987; Maillard et al„ 1992) and

Figure 6b gives an isopach map of the Cenozoic

strata. The structure map illustrates the broadly

saucer shaped conliguration of the Cenozoic

Valencia Trough. The isopach map, in combination

with the bathymetric map, shows that this basin is

clearly partly sediment starved.

Upper and Middle Crust

The available geophysical data permit to

image the present day crustal conliguration ol the

Trough (Figs. 7, 8 and 9). Reflection and refraction

seismic data show that along both margins the

upper-middle crust, almost reflection free, varies in

thickness between 6-10 km and has velocities ot

6. 1-6.2 km/s (Figs. 7 and 9). The overlaying pre¬

rift Mesozoic carbonates are characterized by

velocities ranging between 5.4 to 5.9 km/s (Pascal

et al„ 1992; Torne et al„ 1992; Daiiobeitia et al„

1992). Across the axis of the Trough, COP and

CDP data allow to trace two different scenarios. In

the southern region, seismic data show a series ot

NW dipping reflectors at upper/middle crustal lev¬

els; these are attributed to a Mesozoic basin (e.g.

Mauffret et al., 1992; Torne et al„ 1992) which

may represent the southeastern prolongation of the

Iberian Chain. The thickness of the upper/middle

crust, without Cenozoic sediments, is in this area

of the order of 7-8 km (Fig. 8a and P-1II of Fig. 9).

In the central parts of the Trough there is no evi¬

dence for the presence of thick Mesozoic series

(Fig. 8b).

Seismic refraction/wide-angle reflection data

also reveal that the upper/middle crust thins from

the southwestern off-shore regions in a northeaster¬

ly direction towards the central areas of the Trough

(P-III of Fig. 9). In contrast, profiles P-I and P-II

(Fig. 9) reveal that the thickness of the upper/mid¬

dle crust does not vary significantly across the

Trough; this is in accordance with the ESP data

(e.g., Pascal et al., 1992).

Lower Crust

Reflection seismic data show that beneath the

Ebro Platform the 6-7 km thick lower crust is char¬

acterized by good reflectivity, involving 1-4 km

long subhorizontal reflectors (Fig. 7a; Torne et al.,

1992; Collier et al., 1994). In contrast, the 9-10 km

thick lower crust of the Mallorca margin (P-I of

Fig. 9) is variably reflective, showing disrupted

reflectors which are difficult to trace along the pro¬

file (Collier et al., 1994; Fig. 7b). Considering the

shallow water and lack of near-surface low-veloci¬

ty layers, Collier et al. (1994) favour that disrup¬

tion of the lower crustal reflectors is genuine and

not caused by multiple interference. Towards the

axis of the Trough, the reflective lower crust thins

very rapidly where it appears to be missing (Torne

et al., 1992) or to be reduced to a 1-2 km thick

layer (P-I of Fig. 9), and displays a moderate

velocity gradient of 0.1 s~ ' (Daiiobeitia et al.,

1992). These results are confirmed by coincident

steep and wide-angle reflection profiles, which

show that lower crustal reflectivity diminishes

below the slope-break and vanishes towards the

central parts of the Trough (Gallart et al., 1995;

Source :

Depth to the base of Neogene (m) \ Neogene sediment thickness (m)

PERI-TETHYS MEMOIR 2: ALPINE BASINS AND FORELANDS

113

Source : MNHN, Paris

FIG. 6. a) Dcpih to the base ol Neogene, contour interval 500 m. b) Neogenc sediment thickness, contour interval

500 m. Compiled from Maillard et al. (1992) and Lanaja (1987). Abbreviations as in Fig. 2.

TWT (s) TWT (s)

114

M. TORNE. E. BANDA & M. FERNANDEZ: VALENCIA TROUGH. SPAIN

0

1

2

3

4

5

6

7

8

9

10

a)

0 20 40 60

Distance (km)

FIG. 7. a) Unmigrated CDP line 819 with velocity solution from ESP-6,

b) Unmigrated COP line 822 with velocity solution from ESP-3. No vertical exaggeration

at 4.1 km/s. Inset shows location of seismic sections. Triangles give mid-point locations

of ESP-6 and 3. Abbreviations as in Fig. 2. COP profiles from Collier et al. (1994). ESP

results from Pascal et al. ( 1992) and Torne el al. (1992).

Source : MNHN, Paris

PERI-TETHYS MEMOIR 2: ALPINE BASINS AND FORELANDS

115

>

P

N

vp(km/s)

0 4 6

s

FIG. 8. a) Unmigrated COP line 818 with velocity solution from ESP-7.

b) Unmigrated COP line 821 with velocities solutions from ESP-3. 4 and

5.

c) Unmigrated COP line 808 with velocity solution from ESP-2. No verti¬

cal exaggeration at 4.1 km/s. Me: Messinian; Mz: Mesozoic; UC: Upper

Crust; LC: Lower Crust; M: Moho. Inset shows location of seismic sec¬

tions. Triangles give mid-point locations of ESPs. Abbreviations as in

Fig. 2. COP profiles from Collier et al. (1994). ESP results from Pascal et

al. (1992) and Torne et al. (1992).

Source : MNHN. Paris

116

M. TORNE. E. BANDA & M. FERNANDEZ: VALENCIA TROUGH. SPAIN

Vidal et al., 1995). In the southern parts of the

basin, the lower crust is 5 km thick and is charac¬

terized by average velocities of 6.8 km/s (Pascal et

al., 1992).

Watts et al. (1990) argue that the highly reflec¬

tive lower crustal layer predates the mid-Tertiary

extensional event; as such, underplating processes,

related to the Cenozoic extension, could not cause

the observed reflectivity of the lower crust. This

view is supported by Collier et al. (1994), who

based on the analysis of the lower crustal reflectiv¬

ity patterns in different parts of the Trough, deter¬

mined that Cenozoic extension significantly

weakened or even destroyed the lower crustal

reflectivity.

Moho Topography

In the area of the Valencia Trough, the

crust/mantle boundary is reasonably well con¬

strained by seismic and gravity data. Figure 10

gives a smoothed depth map of the Moho which is

based on an integration of seismic data and gravity

modelling results. The Moho raises gradually from

a depth of 27-30 km under Iberia to 15-16 km

beneath the Trough axis and descends toward the

Baleares to a depth of 22-24 km. The Moho shal¬

lows from 18-19 km at the southwestern end of the

basin to about 13-14 km at its northeasternmost

end where it opens into the Provencal Basin. In

cross-section, the Trough is slightly asymmetric,

having a steeper flank towards the Baleares

(Fig. 10).

In off-shore areas, velocities of the uppermost

lithospheric mantle range from 7.7 to 7.9 km/s,

whereas on-shore Iberia, velocities of 8.1 km/s are

recorded. The upper-mantle velocity increases

from 7. 7-7. 8 km/s beneath the axial parts of the

Trough to 7. 9-8.0 km/s towards the mainland,

whereas beneath the Balearic Promontory, they

remain in the 7. 7-7. 8 km/s range.

On COP the Moho can be traced throughout

the basin (Figs. 7 and 8; Collier et al.. 1994). The

reflection signature of the Moho varies beneath the

different parts of the basin and can be correlated

with differences in the amount of stretching. Ceno¬

zoic extension may have modified the reflection

character of the Moho.

Lithosphere-Asthenosphere Boundary

Information on the configuration of the lithos-

phere/asthenosphere boundary comes primarily

from 2D modelling of the central parts of the

Trough, using gravity and geoid anomalies along a

profile extending from the Ebro basin to the South-

Balearic basin (Watts and Torne, 1992a; Zeyen and

Fernandez, 1994). Watts and Tome (1992a) pointed

out that the two margins of the Trough are charac¬

terized by different structural styles; the Spanish

margin is a rift-type margin whereas the Balearic

margin appears to be a constructional margin

which is underlain by a broad region of lithospher¬

ic thinning. Zeyen and Fernandez (1994) conclude

that the shallowness of the lithosphere/asthenos-

phere boundary indicates that the causal rifting

event had terminated only very recently, as previ¬

ously suggested by Morgan and Fernandez (1992).

Surface-wave and tomography studies (Marillier

and Mueller, 1985; Spakman, 1990) show that the

Trough lies in a region of lithospheric thinning and

anomalous low-velocity sublithospheric upper-

mantle, characteristic for the western Mediter¬

ranean.

BASIN MODELLING

Subsidence Analysis

In an effort to isolate tectonic subsidence from

the effects of sedimentary loading during the evo¬

lution of the Valencia Trough, backstripping analy¬

ses were carried out by several authors (Watts et

al., 1990; Bartrina et al., 1992; Roca and

Desegaulx, 1992) on the basis of wells located on

the western margin of the Trough. The observed

tectonic subsidence curves show an exponential

Source :

PERI-TETHYS MEMOIR 2: ALPINE BASINS AND FORELANDS

117

w

Iberia ▼

P-ll

Mallorca

E

P-lll

SW Iberia y NE

42N

40 N

FIG. 9. Velocity-depth crustal models along profiles P-I. P-II. and P-III. Stippled

zones indicate Neogcne sediments. Dashed lines mark the transition between the

upper and lower crust. Inclined stripped pattern denotes a gradient of 0.1 s' in the

lower crust. Triangles indicate shore-line location. Values show P-wave velocities

in km/s. Inset gives location of profiles (after Danobeitia et al.. 1992).

Source : MNHN . Paris

118

M. TORNE. E. BANDA & M. FERNANDEZ: VALENCIA TROUGH. SPAIN

Depth to the Moho (km)

O’ 2'E

FIG. 10. Smoothed depth to the Moho map. contour interval I km. Compiled

from results of Banda ct al. (1980). Danobeitia et al. (1992). Gallarl et al., (1990),

Pascal ct al. (1992). Tome et al. (1992) and Zeyen et al. (1985).

Source : MNHN. Paris

PERI-TETHYS MEMOIR 2: ALPINE BASINS AND FORELANDS

119

trend with total tectonic subsidence values in the

range of 1.4 to 1.6 km. ID subsidence results also

show that it is difficult to separate the late

Oligocene-early Miocene syn-rift from the post-rift

subsidence (Watts et al., 1990; Roca and

Desegaulx, 1992). The shape of the subsidence

curves must be carefully analyzed, as inaccuracies

in palaeo-waterdepth estimates result in significant

uncertainties (e.g, Watts et al., 1990; Bartrina et al.,

1992) .

Spatial variations in the tectonic subsidence

along a regional transect considering CDP line 821

were analyzed by Watts and Torne (1992a) and

Torres et al. (1993) who concluded that the form of

the backstripped curves is similar, irrespective of

the elastic thickness (Te), providing that Te values

of 5 and 25 km and that corresponding to the

450°C isotherm, are assumed. Such quantitative

subsidence analyses confirm the regional broad

subsidence of the Valencia Trough and its accentu¬

ation to the northeast towards the continent-ocean

transition (Fig. 11). Flexural subsidence analyses

indicate uplift of the Catalan Coastal Ranges and

the Balearic Islands, contemporary with subsidence

with the Valencia Trough. Detailed studies, incor¬

porating fine stratigraphic and palaeo-waterdepths

analyses, are able to distinguish an initial phase of

rapid subsidence (30-15 Ma) followed by a post¬

rift phase, characterized by lower subsidence rates

(Roca and Desegaulx, 1992; Torres and Bois,

1993) .

A 3D backstripping analysis of the entire

region, carried out by Watts and Torne (1992b)

using seismic and well data, confirmed that the

Trough corresponds to a broad region of subsi¬

dence which is flanked by uplift in the Catalan

Coastal Ranges and in the Balearic Promontory.

Maximum tectonic subsidence values of 4.2-

4.4 km are reached in the NE at the continent-

ocean transition (e.g., Pascal et al., 1992), whereas

elsewhere tectonic subsidence values range

between 1 and 3 km (Fig. 1 1).

Numerical Modelling

Numerical models applied to the study area

are far from complete and self-consistent. Never¬

theless, they provide valuable information on the

evolution of the Valencia Trough.

Two key features of this basin are difficult to

handle with "classical” extensional models, namely

the observed crustal and lithospheric asymmetry

across and along its axis, and the fact that the ini¬

tial rifting phase was coeval with compression

along the Balearic margin. The majority of the

numerical approaches, so far performed, omitted

both facts and thus result in the assumption of sim¬

plified models. The difficulties encountered in

modelling are particularly acute for the southern

region of the Trough, where also the surface heat-

flow values are consistently higher and the tectonic

subsidence less than expected (Watts and Torne,

1992a). Therefore, numerical approaches were

applied to the central parts of the Trough and con¬

centrated on determining the pre-rift lithospheric

conditions, the duration of the rifting stage, and the

amount of lithospheric stretching.

Keeping in mind that the area of the Valencia

Trough was uplifted and subjected to erosion dur¬

ing the Paleogene, Morgan and Fernandez (1992)

attempted to evaluate the lithospheric conditions

which prevailed in the western and central regions

prior to the Oligocene-Miocene rifting. A formula¬

tion of lithospheric buoyancy was used to back-

calculate the possible pre-extension lithospheric

structure, consistent with mid-Tertiary elevation

and lithospheric strength constraints. This ID

approach was applied to different lithospheric

columns along a profile from the Ebro basin to the

centre of the Valencia Trough. For the Spanish

margin, stretching factors ranging from 1.45 to

1.87 and an initial crustal thickness between 27

and 35 km were arrived at, whereas for the centre

of the Trough these values are 2.9-4 and 26-36 km,

respectively. The estimated stretching factors can

be reduced by about 10%, assuming erosion of

1 km of pre-rift sediments. Although Morgan and

Fernandez (1992) proposed differential stretching

for the Valencia Trough, where 8erust>®mantle >n

the centre of the Trough but ^Crust<^mantle undcr

the Spanish margin, uniform stretching in the cen¬

tre of the Trough can also Fit the model, provided

an erosion of about 1.5 km of pre-rift sediments is

acceptable.

A second approach, based on a 1 D pure-shear

stretching model, was explored by Foucher et al.

(1992) using as constraints measured heat-flow

120

M. TORNE, E. BANDA & M. FERNANDEZ: VALENCIA TROUGH, SPAIN

4000

3500

3000

sot; ' 2500

2000

v-

FIG. I I . Tectonic subsidence/uplift map, contour interval 5(K) m. Based on llcxural

backstripping of sediment thicknesses given in Fig. 6b. Parameters used: elastic thick¬

ness (Te) of 5 km. water, sediment and mantle densities of 1030, 2400 and

3300 kg/m3, respectively. Modified after Watts and Torne (1992b).

Source : MNHN, Paris

PERI-TETH YS MEMOIR 2: ALPINE BASINS AND FORELANDS

121

data and bathymetry along three transects located

in the central parts of the Trough (for location see

Fig. 2b). Their results show that the northeastern

half of the Valencia Trough could be successfully

explained by a single rifting event lasting from 28

to 22 Ma with crustal stretching factors of 3.5 and

3.3 for the northern transects. However, this

approach fails to explain the relatively high heat

flow and shallow water depths observed in the

southern part of the Trough (transect T4 of

Fig. 2b), even when a multi-stage Neogene rifting

event with different initial crustal and lithospheric

thicknesses is considered. Foucher and co-workers

speculate that this may be the result of a recent

Pliocene-Quaternary event which would favour a

southward propagation of the rift activity.

Fernandez et al. (1995) attempted to model the

southern Valencia Trough, applying a ID uniform

stretching model constrained by thermal and

palaeotherma! data. Their model takes into account

the thermal effects of sedimentation/erosion, com¬

paction, mineral composition and heat production.

This model suggests that the SW Valencia Trough

underwent a complex geodynamic history, includ¬

ing three Mesozoic rifting events, a Paleogene

compressional and an uplift phase, causing erosion

of about 5 km of Late Jurassic and Cretaceous sed¬

iments, and a late Oligocene-early Miocene exten-

sional phase.

Several 2D numerical models were applied

along different transects located in the central

region of the Valencia Trough. These models,

based on a kinematic approach, are mainly con¬

strained by the basement structure, crustal and

mantle stretching factors and surface heat-flow

data. The first model, presented by Watts and

Torne (1992a), was performed along a transect

extending from the Ebro basin to the South

Balearic basin, coinciding off-shore with CDP line

821. The total tectonic subsidence obtained from a

2D backstripping analysis and the measured heat

How were compared with that obtained from a 2D

non-uniform finite stretching model. The best fit

shows a homogeneous stretching with a b factor

increasing from 1.4 beneath the Spanish margin to

3.0 at the centre of the Trough. However, on-shore,

below the Catalan Coastal Ranges, the observed

tectonic uplift requires a broad region of mantle

stretching (Bmant|e= 1 .4) extending well beyond

the region of crustal stretching. It was assumed that

the rifting episode started 24 Ma ago and had a

duration between 8 and 16 Ma, and that the initial

crustal and lithospheric thickness were 31.2 and

125 km, respectively. A combined geoid and gravi¬

ty model, where a temperature-dependent mantle

density is considered, was used to better constrain

the present-day geometry of the lithosphere below

the Balearic margin. The model results reveal

asymmetry at crustal and subcrustal levels between

the flanks of the Trough and, that the region of

lithospheric thinning is much broader than the

width of the initial rift and extends far beyond the

Balearic Promontory.

Torres et al. (1993) modelled the off-shore

areas of the Trough along CDP line 821, using a

2D uniform stretching model, considering local

isostasy, and the loading and thermal effects of

sediments. Unlike Watts and Torne (1992a), it was

assumed that the Valencia Trough underwent a first

rifting event between 25.2 to 20.0 Ma, which had

minor effects on the deep basin and affected most¬

ly the Spanish margin, an that a second rifting

event between 15.2-10.2 Ma, coeval with the open¬

ing of the Tyrrhenian basin, was responsible for the

present-day Moho uplift. However, there is no

clear reflection-seismic evidence for this second

rifting event. These authors also conclude that, in

contrast to the interpretation given by Fontbotc et

al. (1989) and Roca and Desegaulx (1992), the post

mid-Miocene development of the basin does not

conform to a flexural foreland-type mechanism.

Finally, Jansen et al. (1993) proposed a 2D

necking model along two profiles crossing the cen¬

tral part of the Valencia Trough. Basement subsi¬

dence, Moho depth and gravity anomalies were

calculated for different elastic thicknesses, duration

of rifting and necking depths. Results support a

rifting model with a low elastic thickness (5-

20 km), an intermediate depth of necking (17-

33 km) and a finite duration of the stretching event

(16 m.y.), propagating southwestward. Rifting

would have started 24 Ma ago in the northeastern

profile (Mallorca-Barcelona) and 20 Ma ago in the

southwestern profile (Mallorca-Ebro Delta). A

Pliocene uplift, caused by additional mantle thin¬

ning, is proposed to fit the present-day Moho depth

and uplift in the Valencia Trough and the Iberian-

Mediterranean margin. Whether this uplift is relat¬

ed with the Plio-Quaternary extensional phase is

not evident.

122

M. TORNE, E. BANDA & M. FERNANDEZ: VALENCIA TROUGH. SPAIN

DISCUSSION

The Valencia Trough is one of a several exten-

sional basins located in the western Mediterranean

region, which developed under a compressional

regime. It differs, however, from other rifted basins

of the western Mediterranean, such as the Tyrrhen¬

ian and South Balearic basins, in that extension has

not progressed to crustal separation. Available geo¬

logical and geophysical data provide a reasonably

well constrained image of its present-day crustal

structure. It is now well accepted that the Valencia

Trough is characterized by a strongly attenuated

continental crust, underlain by an anomalous low-

velocity upper mantle, and that its lithospheric

structure shows asymmetry both at crustal and sub-

crustal levels. There is still some debate, however,

among the different authors on the cause and mag¬

nitude of extension and on the evolution of the

Trough.

A major problem in understanding the evolu¬

tion of the basin is the poor knowledge on the

structure of the lithosphere and configuration of

the crust before the late Oligocene extensional

phase. It is well established that during the Meso¬

zoic break-up of Pangea the area underwent repeat¬

ed extensional events and that, during Late

Cretaceous and Paleogene times, intraplate com¬

pressional stresses caused inversion of most of the

extensional Mesozoic basins. However, whether or

not the crust recovered its pre-Mesozoic thickness

as a consequence of basin inversion is still a matter

of debate. Therefore, the first question that has to

be solved concerns the effect of Mesozoic exten¬

sion and Paleogene compression on the lithospher¬

ic configuration of the Valencia Trough, prior to

the onset of Oligocene-Miocene extension.

The Oligocene to Recent evolution of the

Trough is characterized by a first rifting stage (late

Oligocene-early Miocene) which can be explained

either by a back-arc model (the Valencia Trough

evolved in a back-arc position relative to the

Kabylian-Calabrian arc) or by southward propaga¬

tion of the European rift system across the West-

Mediterranean fold belts. During the early-middle

Miocene, Balearic thrusting counteracted exten¬

sional forces and aborted rifting. After locking of

the Balearic thrust front, extension resumed during

the latest Miocene and persisted into the Pleis¬

tocene.

Assuming that the Trough is indeed a rift-type

basin and forms part of the Cenozoic rift system of

Western and Central Europe, there are still several

questions that need to be addressed, such as the ini¬

tiation, duration and mechanism of the Oligo-early

Miocene rifting event, the amount of stretching,

and rift-related processes such as lower crustal

reflectivity, underplating, Moho rejuvenation, ther¬

mal regime, etc.

Timing and Duration of Rifting

Surface geological and reflection seismic data

show that crustal extension commenced during

Chattian(?)-Aquitanian times with the deposition

of the first syn-rift sediments and the development

of a block-faulted relief along the western flank of

the Valencia Trough. To the South, however, the

first syn-rift sediments record an upper-Langhian

age. This suggests that the rifting propagated from

North to South, or that the southern areas had a

higher topographic relief when rift activity com¬

menced. Geological data also show that in the

northern and central parts of the Trough the rifting

phase persisted into Langhian-Serravallian times

(Roca and Desegaulx, 1992). This indicates that

the main rifting activity lasted 8-10 Ma. Subsi¬

dence analysis do not help too much in deciphering

the duration of rifting since, as mentioned above, it

is difficult to separate the Oligocene-early Miocene

syn-rift from the post-rift phase. This underlies the

proposal of different durations of the rifting stage,

ranging from 6 m.y. (Foucher et al., 1992) to 8-

16 m.y. (Watts et al., 1992a) and even to multi¬

stage Neogene events (Torres et al., 1993).

Mechanisms of Rifting

Numerical models attempt to establish the

mechanisms of lithospheric extension. Two end-

member models are available, the pure-shear

(McKenzie, 1978) and the simple-shear model

Source : MNHN . Paris

PERI-TETHYS MEMOIR 2: ALPINE BASINS AND FORELANDS

123

(Wernicke, 1985). In the study area, most of the

available geophysical data favours a pure-shear

mechanism, since there is no clear evidence in sup¬

port of a simple-shear model. This does not rule

out, however, that simple shear deformation is

restricted to the upper/middle crust and is com¬

bined with pure shear deformation at deeper levels

(Lister et al., 1991), though such a model has not

yet been tested. Simple shearing would explain

some major upper crustal structural features, such

as low angle faults and fault-block rotations

observed in the upper/brittle part of the crust.

Although the pure-shear model has been success¬

fully applied to off-shore areas, a differential

stretching mechanism is required to fit the uplift of

on-shore areas. In this context, Cabal and Fernan¬

dez (1995), based on thermal evidences from the

easternmost Ebro basin, proposed that during mid¬

dle Miocene to Recent times thinning of the mantle

lithosphere occurred beyond the zone of actual rift¬

ing and extended some distance to the West of the

Catalan Coastal Ranges.

Stretching Values

The amount of Oligocene and younger lithos¬

pheric stretching was obtained mainly from the

crustal configuration and subsidence studies.

Assuming a pre-rift crustal thickness of 30-35 km,

crustal stretching factors, ranging from 1.4-1.55 for

the western flank of the Trough and 3.2±0.25 for

central areas, and upper-crustal thinning ratios of

up to 2.1 were obtained. This value is much larger

than upper-crustal extension ratios deduced from

structural analysis which give values of up to 1.4-

1.5 (Roca and Guimera, 1992). The observed dif¬

ferences between extension and thinning ratios

may be explained by assuming that the thinned

crust is partially inherited from the Mesozoic rift¬

ing events. On the other hand, Morgan and Fernan¬

dez (1992) suggest that upper crustal stretching

factors in the 1.8 to 2.55 range, can be expected if

a mid-Tertiary topographic relief between 0 and

500 m is assumed. Finally, it can not be excluded

that the observed discrepancies between upper

crustal extension by faulting and mid to lower

crustal attenuation involved a Neogene destabiliza¬

tion of the Moho and its upwards displacement by

magmatic delamination of the lower crust (Ziegler,

1992, 1995).

Rift Related Processes

Collier et al. (1994) established a correlation

between the lower crustal/Moho reflectivity and

the degree of crustal attenuation and concluded

that extension could significantly weaken or even

destroy the reflectivity of the lower crust (see also

Watts et al., 1990), but enhanced the reflectivity of

the Moho. Where the lower crust is very thin, its

reflectivity and the Moho reflector are weak or non

existent. In the centre of the Trough, the absence of

a reflective lower crustal layer and the presence of

a single lower crustal reflector (reflector X), which

coincides with a P-wave velocity increase from 6.4

to 7.8 km/s, has led to Collier et al. (1994) to pro¬

pose that this reflector represents the top of a

crustal transition zone, composed of crustal and

mantle material. Further evidence for intrusion of

mantle material into the base of the crust is the

observed high-velocity gradient in the lower

crustal layer (Danobeitia et al., 1992). The occur¬

rence of a crust-mantle transition zone was already

postulated by Banda et al. (1992) in an effort to

explain the observed difference between upper and

lower crustal attenuation and the presence of

anomalous upper mantle velocities recognized

throughout the basin, the thickness of which is of

the order of 20 km, as deduced from gravity mod¬

elling (Torne and Banda, 1988).

The thermal regime of the Trough, as summa¬

rized in Fig. 2b, is not too well constrained by

measurements which show considerable variations

in the central regions of the Trough. This makes it

difficult to decipher whether this area is character¬

ized by intermediate or high heat flow values.

Moreover, the background heat flow of the

Balearic Promontory is quite similar to that of the

Trough axis, except on Menorca where a slight

increase is recorded. This regional heat flow pat¬

tern is in accord with modelling results which

show that the region of lithospheric thinning

extends beneath the Balearic Promontory (Watts

and Tome, 1992a; Zeyen and Fernandez, 1994); as

124

M. TORNE, E. BANDA & M. FERNANDEZ: VALENCIA TROUGH. SPAIN

such, the latter can not be regarded as the conju¬

gate margin of the Iberian margin.

CONCLUSIONS

Although basin modelling has been quite suc¬

cessful in imaging some aspects of the central parts

of the Trough, it was less successful on a basin

wide scale and particularly for the southern areas.

The underlying reason is, that kinematic models

applied do not account for the complex tectonic

evolution of the area. This concerns mainly the

alternation and nearly contemporaneous extension

and compression and the possible southward rift

propagation. Some of these aspects have been

explained in a step-like manner rather than incor¬

porating all of them into a self-consistent model.

Advective heat transport, radiogenic heat produc¬

tion, melt generation and sediment thermal blan¬

keting are not contemplated in most of the models,

yet they may be important factors in controlling the

evolution of the basin.

On the other hand, lithospheric deformations

imposed on these kinematic models are not con¬

strained by constitutive equations. Depending on

the initial lithospheric structure, the strain rates

obtained from stretching factors and the duration

of the rift stage are at odds with the thermo¬

mechanical behaviour of lithosphere. As pointed

out by Negredo et al. (1995b), thermo-mechanical

constraints on kinematic models may result in dif¬

ferent styles of rifting but can also invalidate a pre¬

conceived mode of deformation. Fully dynamic

models are more self-consistent since deformation

is calculated by coupling constitutive and thermal

equations (e.g. Bassi et al., 1993), yet the high non¬

linearity of the equation renders the results very

sensitive to the initial conditions, making it diffi¬

cult to reproduce the present-day crustal structure

of a given extended area.

Therefore, we conclude that current basin

modelling techniques do not allow to properly

account for the evolution of the Valencia Trough.

In addition, even assuming that these techniques

are capable to handle the main tectonic aspects.

there are still some gaps in our knowledge of the

study area such as:

- Is the Valencia Trough an extensional fea¬

ture on its own? An important aspect when mod¬

elling the Trough is whether it should be

considered as an '‘extensional” feature on its own,

or should be considered as a "branch” of a much

wider extensional region covering the western

Mediterranean. If the latter is correct, then the

Trough would form part of the Gulf of Lyons, Cor-

sica-Sardinia rift system which was active prior to

the late Aquitanian-early Burdigalian opening of

the Algero-Proven^al Basin (Ziegler, 1992).

- What was the pre-rift lithospheric confi¬

guration of the Trough? It is not clear whether or

not the Paleogene compressional phase was large

enough to restore the thinned Mesozoic lithosphere

to its initial conditions. In other words, is the pre¬

sent-day lithospheric structure mainly the result of

the Neogene rifting? or is it partially inherited from

Mesozoic extension?. The Alpine deformation of

the area occupied by the Valencia Trough is still

poorly understood. Therefore, interactive forward

and backward models must be developed which

take into account various degrees of Mesozoic

extension and Paleogene inversion and pre-rift ero¬

sion.

- Is the present-day lithospheric structure

the result of a finite rifting episode or a multi¬

stage rifting? As pointed out earlier, subsidence

analysis do not permit to clearly separate syn-rift

and post-rift phases; correspondingly different

durations of rift activity have been proposed. This

is mainly due to the lack of knowledge of several

parameters, which strongly influence backstripping

results, such as palaeo-bathymetry, precise timing

and amount of erosion, pre-Neogene basement

morphology and density variations. The superposi¬

tion of a thrust loading phase on the southeastern

parts of the Trough is an additional complication

factor. In this respect, stress-induced lithospheric

deflections may also have played an important role

(Cloetingh and Kooi. 1992).

- What is the configuration of the litho¬

spheric mantle and the lithosphere-astheno-

sphere boundary? Information on the structure ot

the lithospheric mantle and the topography of the

base of the lithosphere comes only from thermal

modelling and surface-wave studies. There is no

information on the deep structure of the lithosphere

Source : MNHN, Paris

PERI-TETHYS MEMOIR 2: ALPINE BASINS AND FORELANDS

125

along ihe basin axis and its transition to the ocean¬

ic Algero-Proven9al basin. The lateral extent and

thickness of the low-velocity anomalous upper

mantle is unknown. Passive and active source seis¬

mology in the form of deep seismic soundings and

intermediate tomographic inversion schemes may

help to outline the topography of the base of the

lithosphere and to image the shape of the low-

velocity upper mantle layer.

- Was the Moho a passive marker through¬

out Cenozoic extension? There is some evidence

that the geophysically defined crustymantle bound¬

ary was destabilized during Oligocene and younger

rifting. The differences observed between the mag¬

nitude of crustal extension by faulting and crustal

thinning could at least partly be explained by an

upward displacement and rejuvenation of the

Moho. In this respect, the presence of an anom¬

alous low-velocity uppermost mantle or "transi¬

tional” layer could be explained by the presence of

partial melts which may have interacted with the

lower crust. However, whether the top or the bot¬

tom of this anomalous layer correspond to the pre¬

sent crust/mantle boundary, remains to be

investigated.

Acknowledgements- Our work has been sup¬

ported by the Consejo Superior cle Investigaciones

Cientificcis , CSIC ( Spain j, NATO grant #

CRG/890570 and CIRIT grant # GRQ93-8049.

The authors are very grateful to Dr. Peter Ziegler

who spent much of his time revising and improv¬

ing our original manuscript. Thanks are extended

to Profs. Sierd Cloetingh and Frank Horvath for

their constructive revision of this manuscript.

Some figures shown in this paper were produced

using GMT software (Wessel and Smith, 1991).

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