Dynamics of complex intracontinental basins: the Central European Basin System

Dynamics of Complex Intracontinental Basins

Ralf Littke · Ulf Bayer · Dirk Gajewski · Susanne Nelskamp (Eds.)

Dynamics of Complex Intracontinental Basins The Central European Basin System

With CD-ROM

123

Prof. Dr. Ralf Littke RWTH Aachen Lehrst. Geologie, Geochemie u. Lagerst¨atten Erd¨ol u. Kohle Lochnerstr. 4-20 52056 Aachen Germany [email protected]

Prof. Dr. Ulf Bayer GeoForschungsZentrum Potsdam Telegrafenberg 14473 Potsdam Germany [email protected]

Prof. Dr. Dirk Gajewski Universit¨at Hamburg Institut f¨ur Geophysik Bundesstr. 55 20146 Hamburg Germany [email protected]

Susanne Nelskamp RWTH Aachen Lehrst. Geologie, Geochemie u. Lagerst¨atten Erd¨ol u. Kohle Lochnerstr. 4-20 52056 Aachen Germany [email protected]

ISBN: 978-3-540-85084-7

e-ISBN: 978-3-540-85085-4

Library of Congress Control Number: 2008932466 c 2008 Springer-Verlag Berlin Heidelberg  This work is subject to copyright. All rights are reserved, whether the whole or part of the material is concerned, specifically the rights of translation, reprinting, reuse of illustrations, recitation, broadcasting, reproduction on microfilm or in any other way, and storage in data banks. Duplication of this publication or parts thereof is permitted only under the provisions of the German Copyright Law of September 9, 1965, in its current version, and permission for use must always be obtained from Springer. Violations are liable to prosecution under the German Copyright Law. The use of general descriptive names, registered names, trademarks, etc. in this publication does not imply, even in the absence of a specific statement, that such names are exempt from the relevant protective laws and regulations and therefore free for general use. Cover design: deblik, Berlin Typesetting: Camera-ready by Eva Nelskamp and the Editors Printed on acid-free paper 9 8 7 6 5 4 3 2 1 springer.com

Preface

Although sedimentary basins are visually less impressive than mountain belts, they are mankind’s most important archives, not only regarding our understanding of the past but also in terms of economic and ecological aspects. Subsiding over millions of years, they accumulate sediments which allow us to reconstruct long term climatic and tectonic changes that have affected the history of the Earth. During subsidence and consequent burial the deposits are exposed to greater depth, pressure, and temperature. A chemical kitchen develops, reacting as a geo-reactor, generating fluids such as petroleum and natural gas. Beside the fact that basins provide more than 90% of our energy resources, they also provide large reservoirs of drinking water. Other typical resources are sandstones, carbonates, gypsum and different types of salt. The latter not only includes conventional sodium chloride (kitchen salt, halite), but also potassium salts used as fertilizers. Increasingly, sedimentary basins will also act as long term resources for geothermal energy which may partly fill the forthcoming gap in energy resources. Additionally, salt structures and other layers of low permeability in deep sedimentary basins are in the focus for waste disposal, especially radiogenic and toxic agents. Finally, porous and permeable layers overlain by less permeable cap rocks can be used to store liquids or gases, including carbon dioxide, in the attempt to store energy or to construct climate neutral power plants. Usage of basins is becoming more and more competitive in terms of production of resources and long term storage of waste. In order to achieve compatibility, reliable management will be necessary. Reliable management, however, requires a profound knowledge of the structure to be managed and this knowledge must encompass not only its present state but also, as far as possible, its past and future evolution. For simple basins, straight forward analytic and modeling methods have been developed in various disciplines. There is, however, a class of basins that are either rather large and/or have suffered a complicated geodynamic history over a long time for which classical models fail or are at least insufficient to predict details of the internal structure. We call these basins “complex”. One such is the Central European Basin System (CEBS) extending from Norway to Germany and from Great Britain to Poland. The CEBS has experienced various phases of subsidence and uplift during the past 270 million years. This area will be the central focus throughout this book, although other areas of the world will also be discussed. In 1999, ten years after the fall of the Berlin Wall we took the opportunity to launch a project concerning sedimentary basins, focusing on the Central European Basin System within a future orientated program of the German Research Foundation DFG and the Federal Ministry of Education and Research BMBF, (“GEOTECHNOLOGIEN: The System Earth – From Processes to Management”). At the same time, the German hydrocarbon industry announced that they would provide previously classified data for basic research. Finally, in 2002, it was possible to initiate a special research project SPP 1135: “Dynamics of sedimentary basins”, funded by the DFG for six years as a “Priority Programme” (Schwerpunktprogramm) and cosponsored by the DGMK (German Society for Petroleum and Coal Science and Technology) as representative of the German hydro-

VI

Preface

carbon industry (DGMK-project 577). For the following six years basic research was performed in about 30 projects and results were presented in a variety of publications, including three peer-reviewed special volumes (Marotta and Bayer, 2005, Littke et al. 2005, Bayer et al. 2008). The special research project SPP1135 was built on scientific research into the area from the last twenty years. It was first of all the European Geo-Traverse (Blundell et al. 1992) crossing the area in the late 1980s. EUROPROBE served as an umbrella bringing together scientists from eastern and western Europe after the political boundaries became transparent in the 1990s. Sub-programmes like TESZ (Trans-European Suture Zone) and associated programmes like PACE (Paleozoic Amalgamation of Central Europe), DEKORP Basin’96, MonaLisa, Polonaise and Celebration 2000, and Thor need to be mentioned here. All these projects, funded by different sources, provided a sound basis for understanding in particular the deeper crust below the Central European basin system by focussed research. Having published many results of our research program in scientific journals, we felt the need to present our major findings in a book specifically designed to describe a basin system completely and to elucidate the major processes acting therein. This is the first objective of this book, and the sedimentary system selected is that of the Central European Basin System, which is one of the largest and most complex continental basins on Earth. In order to develop a conclusive concept for the structure and evolution of this outstanding example of a complex basin, it was necessary to integrate the data and to organise the people specialised in dealing with certain data sets into a non-hierarchical scientific system, which was a considerable enterprise given that the databases came from geology, geophysics, geochemistry, hydrogeology and so on. The contribution by authorities from different fields of geoscientific research also provided the rare opportunity to combine expert knowledge from different disciplines in one book. In the course of our research programme, we could learn a lot about the disciplines of “the others”, thereby obtaining a wider view of sedimentary basin dynamics. We want to share this experience with our readers and have encouraged the authors to describe basic processes from their view as a geophysicst, sedimentologist, structural geologist, petroleum geochemist, hydro-geochemist and so on. Thus, the second objective of this book is to provide an advanced understanding of some of the most important interpretation concepts and parameters relevant for understanding processes acting in sedimentary basins. In this sense, this book should be regarded as an advanced teaching book bringing together expertise from different scientific disciplines. This expertise can be applied to sedimentary basins in general, not only the Central European Basin System. The result we present here is a multi-authored book, whereby the authors of the different chapters are responsible for the content of their chapters. Finally, we would like to remember with gratitude our colleagues Friedrich Theilen (Kiel) and Hartmut Jödicke (Münster). Acknowledgement This book would not have been possible without the support of the German Research Foundation (DFG) funding of SPP1135, and the German Society for Petroleum and Coal Science and Technology (DGMK) as the representative organisation of the German hydrocarbon industry (DGMK-project 577). Many people from these two institutions and the German petroleum industry have greatly contributed to our project, including the DFG review committee and the industrial representatives of the petroleum companies Wintershall Holding AG, Kassel, RWE Dea AG, Hamburg, Gaz de France Production

Preface

Exploration Deutschland GmbH, Lingen, and ExxonMobil Production Deutschland GmbH, Hannover. We would like to thank all of them, in particular Thilo Bechstaedt, Sören Dürr, Martin Jentsch, and Ingrid Winter for their great support. We also acknowledge the help of the following persons: Benjamin Bruns greatly helped to organize the reference list of this book, Martin Koppelberg and Gabor Lang helped with the key word and abbreviation index, Eva Nelskamp made the proofs, Hilary Horsfield read earlier drafts of the manuscripts and homogenized the usage of the English language as much as possible and Frederik Orth improved the readability of several of our figures. Furthermore, the authors of the individual chapters wish to extend their thanks to the following people and organizations: Chapter 3.2: Derek Blundell for discussion on Caledonian tectonics and EEG Erdöl Erdgas GmbH for providing the Rerik dataset and for permission to publish our results. Chapter 3.3: Nederlandsche Aardolie Maatschappij (NAM; a Shell operated 50/50 joint venture between Shell and ExxonMobil) is thanked for use of the Groningen seismic volume. We thank RWE-DEA A.G. for 3D seismic and borehole data. We thank the Landesamt für Geologie und Bergwesen (LAGB) Sachsen-Anhalt for data and support. Chapter 3.4: Ulrich Glasmacher, Heidelberg, for discussions on the fission track interpretation, Marek Narkiewicz, Warsaw, Andreas Beha, Copenhagen, and Petra David, Utrecht, for valuable information on the Polish, Danish, and Dutch part of the basin, respectively. Chapter 4.1: Dr. G. Beutler, J. Barnasch, M. Franz, Halle, Dr. H. W. Kozur, Budapest, and Dr. F. Kockel, Hannover, for valuable information and discussions; J. Barnasch helped additionally with computer graphics. Dr. E. Brand, Kassel, provided useful unpublished information on the late Bajocian unconformity. Chapter 4.3: Dr. H. W. Kozur, Budapest is thanked for valuable information and discussions. Chapter 4.5: Our neotectonic studies in Schleswig Holstein were strongly supported by Dr. Sven Christensen, head of the Geological Division of the „Landesamt für Natur und Umwelt des Landes Schleswig Holstein“. Chapter 5.3: We thank GDF Produktion Exploration Deutschland, EMPG and EWE Aktiengesellschaft for providing a high-quality data set. Mike Hudec, Martin Jackson and John Warren are thanked for their valuable contributions during this project. Christoph Krämer is thanked for his drafting support. Chapter 6.4: GFZ-Potsdam, BGR, RWTH Aachen and StatoilHydro are thanked for support and permission to publish. Chapter 6.5: Anne Richter, FZ-Jülich, and Yves Gensterblum, RWTH Aachen, RWTH Aachen for valuable technical assistance in the pyrolysis experiments; Anke Jurisch for evaluating experimental data and assistance in preparing the figures.

Ralf Littke

Ulf Bayer

Dirk Gajewski

Susanne Nelskamp

VII

Contents

1

2

3

3.1

3.2

3.3

Characteristics of complex intracontinental sedimentary basins

1

(U. Bayer · H.-J. Brink · D. Gajewski · R. Littke) 1.1 Introduction 1.2 Classifications of basin complexity 1.3 Summary

3 3 12

The Central European Basin System – an Overview

15

(Y. Maystrenko · U. Bayer · H.-J. Brink · R. Littke) 2.1 Introduction 2.2 Crustal association 2.3 Permian Basin formation and subsequent subsidence 2.4 Subsequent formation of sub-basins 2.5 Sedimentary history 2.6 Fluids within the Central European Basin System 2.7 The Central European Basin System – prototype of a complex sedimentary basin

Strain and temperature in space and time

Driving mechanisms for basin formation and evolution (M. Cacace · U. Bayer · A.M. Marotta · C. Lempp) 3.1.1 Driving mechanisms for basin evolution 3.1.2 Kinematic models for basin formation 3.1.3 Rheological models 3.1.4 Modelling complex basins Crustal structures and properties in the Central European Basin System from geophysical evidence (C.M. Krawczyk · W. Rabbel · S. Willert · F. Hese · H.-J. Götze · D. Gajewski & the SPP-Geophysics Group) 3.2.1 Introduction 3.2.2 Structural inventory and physical properties from seismic observations 3.2.3 Conductive layers and bodies from magnetotelluric observations 3.2.4 Rock properties and density structure from potential field investigations 3.2.5 Summary Strain and Stress (J. Kley · H.-J. Franzke · F. Jähne · C. Krawczyk · T. Lohr · K. Reicherter ·M. ScheckWenderoth · J. Sippel · D. Tanner · H. van Gent - the SPP Structural Geology Group) 3.3.1 Introduction 3.3.2 Structural framework of the Central European Basin System 3.3.3 Structural analysis and quantification of strain 3.3.4 Stress history 3.3.5 The Central European Basin Systems structural evolution

17 19 22 25 26 30 34

35 37 37 37 50 66 67

67 68 82 85 94 97

97 102 105 116 121

X

Contents

3.4

4

4.1

4.2

Subsidence, inversion and evolution of the thermal field (R. Littke · M. Scheck-Wenderoth · M.R. Brix · S. Nelskamp) 3.4.1 Introduction 3.4.2 The Central European Basin System as example of regional subsidence models 3.4.3 Temperature in sedimentary basins 3.4.4 Maturity and temperature parameters in sedimentary basins 3.4.5 Variability of palaeotemperature fields in the Central European Basin System

Basin fill

Depositional history and sedimentary cycles in the Central European Basin System (G.H. Bachmann · T. Voigt · U. Bayer · H. von Eynatten · B. Legler · R. Littke) 4.1.1 Palaeoclimate, palaeogeography and palaeoenvironment 4.1.2 Sedimentary cycles 4.1.3 Provenance of sediments in the Central European Basin Basin initiation: Volcanism and sedimentation (Ch. Breitkreuz · M. Geißler · J. Schneider · H. Kiersnowski) 4.2.1. Late Palaeozoic basins in central Europe – distribution, volcanic activity and magmagenetic aspects 4.2.2. Data base, distribution and volumes of Late Palaeozoic volcanics in the Central European Basin System 4.2.3. Stratigraphy and geochronology of volcanic successions in the Southern Permian Basin 4.2.4. Volcanic facies in the Southern Permian Basin 4.2.5. Syn- to postvolcanic sedimentation during the Lower Rotliegend and Upper Rotliegend I 4.2.6. Landscape evolution during the initial phase of the Southern Permian Basin

125 125 125 133 137 141

155 157 157 161 169 173

173 173 175 176 178 179

4.3 Upper Rotliegend to Early Cretaceous basin development (H. Stollhofen · G.H. Bachmann · J. Barnasch · U. Bayer · G. Beutler · M. Franz M. Kästner · B. Legler · J. Mutterlose · D. Radies) 4.3.1. Introduction 4.3.2. Upper Rotliegend II 4.3.3. Zechstein 4.3.4. Buntsandstein 4.3.5. Muschelkalk 4.3.6. Keuper 4.3.7. Jurassic 4.3.8. Early Cretaceous

181

4.4 Sedimentation during basin inversion (T. Voigt · K. Reicherter · H. von Eynatten · R. Littke · S. Voigt · J. Kley) 4.4.1. Introduction 4.4.2. Basin formation 4.4.3. Effects of basin inversion on deposition 4.4.4. Sedimentation during inversion in the Central European Basin 4.4.5. The North German Basin during the Tertiary

211

4.5 Glaciation, salt and the present landscape (F. Sirocko · K. Reicherter · R. Lehné · Ch. Hübscher · J. Winsemann · W. Stackebrandt) 4.5.1. Introduction 4.5.2. Modern topography and glacial isostasy 4.5.3. Crustal movements, seismicity and landscape formation

181 182 185 188 191 194 199 207

211 211 215 220 228 233 233 233 236

Contents

5

5.1

5.2

5.3

5.4

5.5

6

6.1

Salt dynamics

Salt as sediment in the Central European Basin System as seen from a deep time perspective (J.K.Warren) 5.1.1. Introduction 5.1.2. Mother brines: isochemical systems? 5.1.3. Evaporite sediments and climate 5.1.4. Evaporite volumes in deep time 5.1.5. Evaporite volumes & tectonics? 5.1.6. Episodic halokinesis

247 249 249 251 255 259 262 267

Flow and transport properties of salt rocks (J.L. Urai · Z. Schléder · C.J.Spiers · P.A. Kukla) 5.2.1. Introduction 5.2.2. Physical properties of evaporites 5.2.3. Deformation mechanisms and rheology of halite in experiments 5.2.4. Deformation mechanisms and rheology of carnallite and bischofite 5.2.5. Natural laboratories 5.2.6. Discussion and outlook

277

Dynamics of salt structures (P.A. Kukla · J.L. Urai · M. Mohr) 5.3.1. Introduction 5.3.2. Concepts of salt tectonics 5.3.3. Salt geometries and kinematics – a case study 5.3.4. Salt sediment interaction 5.3.5. Multiphase salt dynamics in the Central European Basin System

291

Dynamics of salt basins (M. Scheck-Wenderoth · Y. Maystrenko · C. Hübscher · M. Hansen · S. Mazur) 5.4.1. Introduction 5.4.2. Regional pattern of salt structures in the Central European Basin System 5.4.3. History of salt movements in the Central European Basin System 5.4.4. Case study Glückstadt Graben 5.4.5. Case study NE German Basin 5.4.6. Case study SW Baltic Sea 5.4.7. General findings for salt-containing intra-continental basins

307

Temperature fields, petroleum maturation and fluid flow in the vicinity of salt domes (F. Magri · R. Littke · S. Rodon · U. Bayer · J.L. Urai) 5.5.1. Introduction 5.5.2. Impact of salt structures on temperature field and oil maturation 5.5.3. Fluid flow in salt 5.5.4. Impact of salt structures on groundwater transport processes within sedimentary basins

Fluid systems

Fluids in sedimentary basins: an overview (R. Gaupp · P. Möller · V. Lüders · R. di Primio · R. Littke) 6.1.1. Relevance of geofluids 6.1.2. Definitions 6.1.3. Subsurface aquaeous fluids 6.1.4. Petroleum fluids

277 278 278 284 284 289

291 292 293 301 304

307 308 309 315 318 320 321 323 323 323 328 330 345 347 347 347 348 359

XI

XII

Contents

6.2

6.3

6.4

6.5

Transport processes (J.L. Urai · G. Nover · C. Zwach · R. Ondrak · R. Schöner · B.M. Krooss) 6.2.1. Introduction 6.2.2. Physical mechanisms and concepts 6.2.3. Fault seals and top seals 6.2.4. Geological aspects of fluid transport

367

Fluid-rock interactions (R. Schöner · V. Lüders · R. Ondrak · R. Gaupp · P. Möller) 6.3.1. Introduction 6.3.2. Evolution of deep brines 6.3.3. Palaeo-fluid reconstruction 6.3.4. Organic-inorganic interactions 6.3.5. Modelling fluid-rock interactions 6.3.6. Geological applications

389

Petroleum systems (R. di Primio · B. Cramer · C. Zwach · B.M. Krooss · R. Littke) 6.4.1. Concepts of petroleum system modelling 6.4.2. Petroleum source rocks 6.4.3. Shallow and microbial gas 6.4.4. Sources of deep gas 6.4.5. Petroleum alteration – biodegradation 6.4.6. Overpressured reservoirs 6.4.7. Effects of glaciation on petroleum systems

411

367 367 372 386

389 389 391 401 404 408

411 413 418 422 425 428 430

Origin and distribution of non-hydrocarbon gases (B.M. Krooss · B. Plessen · H.G. Machel · V. Lüders · R. Littke) 6.5.1. Introduction 6.5.2. Nitrogen 6.5.3. Carbon dioxide 6.5.4. Hydrogen sulfide 6.5.5. Evidence from vein mineralisation and fluid inclusions

433

References

459

Subject Index

507

433 433 443 447 457

Abbreviations

= Werra-Anhydrites = Stassfurt-Anhydrites = Altmark-Brandenburg Basin = Avalonia-Baltica Suture = Apatite Fission Track Analysis = Aller Lineament = Apatite = American Petroleum Institute = Allertal = Autun Basin = Athesian Volcanic District = Boskovice Graben = Bornholm-Darlowo Fault Zone = Broad Fourteens Basin = Federal Institute for Geosciences and Natural Resources = British Geological Survey BGS = British Institutions Reflection BIRPS Profiling Syndicate = Blanice Graben BLG = Formation Volume Factor (of oil) Bo = Barrels of Oil Equivalent BOE = Before Present BP = Solid Bitumen Reflectance BRr = North-East Brandenburg Swell BS = Bacterial Sulfate Reduction BSR = Bourbon l´Archambault Basin BU = Bohemian Vindelician Massif BVM = Carpatian Basin CA = Zechstein Limestone Ca1 = Zechstein 2 Carbonate (Stassfurt Ca2 Carbonate) = Conodont Colour CAI = Collio Basin CB = Caledonian Deformation Front CDF = Canyon Diablo Troilite (sulfur CDT istotope standard) = Central European Basin System CEBS = Central Graben CG = Central Glückstadt Graben CGG = Common Mid Point CMP = Central Netherlands Basin CNB = Organic Carbon Corg = Carbon Preference Index CPI = Carnic Alps CR = Common Reflection Surface CRS Cryo-SEM = Cryo-Scanning Electron Microscopy = Cenozoic Cz = Detfurth D = Deep Aquifer Complex DAC = Danish Basin DB A1 A2 ABB ABS AFTA AL Ap. API AT AU AVD BCG BDF BFB BGR

= Donezk Basin = Depositional Cycle = Dielectric Constant Measurement = Deutsches Kontinentales Reflexionsseismisches Programm = Döhlen Basin DÖ = Dowsing–South Hewett Fault Zone DSHFZ EA=EAS = Eichsfeld-Altmark Swell = Erzgebirge Basin EB = Electron Backscatter Diffraction EBSD = East Brandenburg Sub-Province EBSP = East European Craton EEC = East European Platform EEP = Elbe Fault System EFS = Elbe Fault Zone EFZ = Enhanced Gas Recovery EGR = European Geotraverse EGT = Eastholstein Trough EHT = Ems Low EL = Elbe Line (see EOL) EL = Elbe-Odra Line EOL = Enhanced Oil Recovery EOR = Equation of State EOS = Erosion Rate ER = East Sudetic Depression ESD = Emsland Trough ET = Flechtingen-Altmark Sub-Province FASP = Finite Element FE FEFLOW = Finite Element Subsurface FLOW System = Fluid Inclusions FI = Franconian Line FL = Flechtingen-Roßlau Block FL = Flechtingen High Fle = Formation Fm = Franconian Basin FR = Fennoscandian Shield FS = Field Size Distribution FSD = Foresudetic Monocline FSM = Fjerritslev Trough FT = Fission Track FT = Fourier Transform Infrared FTIR Spectroscopy = Flow Zone Indicator FZI GC/IRMS = Gas-Chromatography / Isotope Ratio Mass Spectrometry = Gardelegen Escarpment GE = Geological Survey of Denmark and GEUS Greenland = Glückstadt Graben GG = Glacial Isostatic Adjustment GIA DB DC DCM DEKORP

XIV

Abbreviations GIS GL GMWL GOR GP GPS GR Grt. GSH GSSP GWC H Ha HAH HC HCM HDLC HF HG HI HL HPHT HRF HT ID-TIMS IF InSAR IS IS J K1 K2 KFC KM 1 KM 2 KM 4 KP KS KU La LAB LB LBM LC LE LGM LH LN LO LSB LU Ma MAGT MCR MDH MFB MH MIM ML MM MM MNSH

= Geographic Information System = Gravity Lineament = Global Meteoric Water Line = Gas to Oil Ratio = Guardia Pisano Basin = Global Positioning System = Gamma Ray = Garnet = Grand Sillon Houllier Fracture Zone = Global Stratotype Section and Point = Gas-Water Contact = Hardegsen = Harz = Hamburg High = Hydrocarbon = Holy Cross Mountains = High Density Lower Crust = Hydrofluoric Acid = Horn Graben = Hydrogen Index = Helgoland Low = High Pressure High Temperature = Hunsrück Fracture Zone = Hamburg Trough = Isotope Dilution – Thermal Ionisation Mass Spectrometry = Ilfeld Basin = Interferometric Synthetic Aperture Radar = Iapetus Suture = Intra Sudetic Basin = Jurassic = Upper Cretaceous = Lower Cretaceous = Koszalinseyre-Tornquist-Zone = Grabfeld-Formation = Middle Keuper 2 = Middle Keuper 4 = Krkonoše Piedmont Basin = Kuiavian Segment = Lower Keuper = Lausitz Thrust = Lithosphere-Astenosphere Boundary = Leer-Bremen Fault Zone = London-Brabant Massif = Lu Caparoni Basin = Lausitz Escarpment = Last Glacial Maximum = Lusatian High = Landshut-Neuötting = Lodève Basin = Lower Saxony Basin = Lysogory Unit = Million Years = Mean Annual Ground Temperature = Mid-German Crystalline Rise/Kyffhäuser = Magdeburg-Dessau High = Moary Firth Basin = Moho = Multiple Inverse Method = Mölln Low = Middle Muschelkalk = Malopolska Massif = Mid North Sea High

Mo MO MO MORB MPI MPT MSK MSR MT MU MVSP NASC NBL ND NDA NDB NEGB NGB NHF NKH NL NL NN NPI NPB NS NS Ntot OAE Od OI Os OWC PAAS PAZ PB PB PBF PD PDB pdHg Pf Pf PH PH PI PIG PIXE ppm PS PS PS Psat PT P-T P-T-X PVT PVTt PT R Ra RB

= Bending Moment = Upper Muschelkalk = Montceau les Mines Basin = Mid-Ocean Ridge Basalts = Methylphenantrene Index = Mid-Polish Trough = intensity=Medvedev-SponheuerKarnik scale of seismic intensity = Microbial Sulfate Reduction = Magnetotelluric = Lower Muschelkalk = Mecklenburg-Vorpommern Sub-Province = North American Shale Composite = Neiße-Bobr-Low = North Danish Basin = Numerical Dynamic Analysis = Norwegian-Danish Basin = North East German Basin = North German Basin = Northern Harz Boundary Fault = Norderstedt Kiel High = Netherlands Low = North Netherlands Low = North Netherlands Swell = Normalized Porosity Index = Northern Permian Basin = North Sudetic Basin = Netherlands Swell = Total Nitrogen = Oceanic Anoxic Event = Odra = Oxygen Index = Osning = Oil-Water Contact = Post-Archean Australian Shale = Partial Annealing Zone = Polish Basin = Pompecky Block = Pays de Bray Fracture = Perdasdefogu Basin = PeeDee Belemnite = Mercury-Air Displacement Pressure = Pfahl. Sub-Basins = Fluid-pressure = Pritzwalk High = Pennine High = Production Index = Polish Geological Institute = Proton Induced X-ray Emission = Parts Per Million = Polysulfides = Pompeckj-Swell (Block) = Pomeranian Segment of Mid-Polish Trough = Saturation Pressure = Polish Trough = Pressure-Temperature = Pressure-Temperature-Composition = Pressure-Volume-Temperature = PVT through Time = Polish Trough = Röt = Rayleigh Number = Rheder Moor-Blenhorst

Abbreviations REE REY RFH RG RGL RL RM RQI RS RS RT RÜ RVG RWE AG SACS SB SB SCB Scf/STB SD SH SHB SM Sm3 SMOW SNB SNF SO SOM SP SPB SPBV SPI SR St ST STZ SU T1 T2 T2-3 TAI TB TC TDS TEF

= Rare Earth Elements = Rare Earth Elements and Yttrium = Ringkoebing-Fyn-High = Roer Graben = Rhein Graben Lineament = Rheinsberg Lineament = Rhenish Massif = Reservoir Quality Index = Rügen Swell = Rheic Suture = Rheinsberg Trough = Rügen = Roer Valley Graben = RWE AG=Rheinisch-Westfälisches Elektrizitätswerk AG = Saline Aquifer CO2 Storage = Silverpit Basin = Saale Basin = Subhercynian Cretaceous Basin = Standard Cubic Feet per Stock Tank Barrel = Saale Depression = Sylt High = Subhercynian Basin = Middle Bunter = Standard Cubic Meters = Standard Mean Ocean Water = Saar-Nahe Basin = Sveconorwegian Front = Upper Bunter = Sedimentary Organic Matter = Sole Pit Basin = Southern Permian Basin = Southern Permian Basin Volcanic Zone = Source Potential Index = Sedimentation Rate = Steinhuder Meer = St. Etienne Basin = Sorgenfrei-Tornquist Zone = Lower Bunter = Buntsandstein = MiddleTriassic; Lower Triassic = Upper Triassic = Spore Colour = Trier Embayment = Top Cretaceous Reflector = Total Dissolved Solids = Transeuropean Fault

TF THZ Ti Tmax TNO TOC TOR Tpeak TP TS TSR TTI TTZ TW TW TWT UCS UK USA USGS V VDF VF VH VM Vo VRr WBH WBT WCB WD We WEI WHT WNB WSM WSMP WSP Wt.-% Z Z1 ZÖ ZPAZ

= Thuringian Forest Basin = Thrust Zone = Titane = Temperature of Maximum Pyrolysis Yield = Netherlands Organisation for Applied Scientific Research = Total Organic Carbon = Teleseismic Tomography across the Tornquist zone = Maxiumum Paleotemperature = Top Palaeozoic Reflector = Thor Suture = Thermochemical Sulfate Reduction = Time Temperature Index = Tornquist-Teisseyre Zone = Thüringer Wald = Thuriningian-West Brandenburg Depression = Two Way Traveltime = Unconfined Compressive Strength = United Kingdom = United States of America = United States Geological Survey = Volpriehausen = Variscan Deformation Front = Variscan Front = Violet Horizons = Volatile Matter Yield = Vertical Line Load = Mean Vitrinite Reflectance = West Brandenburg High = West Brandenburg Trough = Western and Central Bohemian Basins = Weser Depression = Weser Trough = Weissig Basin = Westholstein Trough = West Netherlands Basin = World Stress Map = World Stress Map Project = West Schleswig Platform = Weight-% = Base Zechstein Reflector = Zechstein = Zöbingen = Zircon Partial Annealing Zone



XV

Authors

Gerhard H. Bachmann Martin Luther Universität Halle-Wittenberg Institut für Geowissenschaften Von-Seckendorff-Platz 3 · 06120 Halle, Germany [email protected]

Filiz Bilgili Christian-Albrechts-Universität zu Kiel Institut für Geowissenschaften, Geophysik Otto-Hahn-Platz 1 · 24118 Kiel, Germany [email protected]

Jens Barnasch Martin-Luther-Universität Halle-Wittenberg Institut für Geowissenschaften Von-Seckendorff-Platz 3 06120 Halle, Germnay [email protected]

Chistoph Breitkreuz Technische Universität Bergakademie Freiberg Institut für Geologie und Paläontologie Bernhard-von-Cotta-Straße 2 09599 Freiberg, Germany [email protected]

Ulf Bayer GeoForschungsZentrum Potsdam Sektion 4.3 Telegrafenberg, C425 · 14473 Potsdam, Germany [email protected]

Heinz-Jürgen Brink Hindenburgstraße 39 · 30175 Hannover, Germany [email protected]

Mikhail Baykulov Universität Hamburg Institut für Geophysik Bundesstraße 55 · 20146 Hamburg, Germany [email protected] Gerhard Beutler Martin-Luther-Universität Halle-Wittenberg Institut für Geowissenschaften Von-Seckendorff-Platz 3 · 06120 Halle, Germany

Manfred R. Brix Ruhr-Universität Bochum Institut für Geologie, Mineralogie und Geophysik 44780 Bochum, Germany [email protected] Holger Busche Geozentrum Hannover Bundesanstalt für Geowissenschaften und Rohstoffe (BGR) Stilleweg 2 · 30655 Hannover [email protected]

XVIII Authors Mauro Cacace GeoForschungsZentrum Potsdam · Section 4.3 Telegrafenberg, C427 · 14473 Potsdam, Germany [email protected]

Martin Bak Hansen Hydro Oil & Energy Research Centre Bergen P.O. Box 7190 · N-5020 Bergen, Norway [email protected]

Bernhard Cramer Geozentrum Hannover Bundesanstalt für Geowissenschaften und Rohstoffe (BGR) Stilleweg 2 · 30655 Hannover [email protected]

Laska Hengesbach Westfälische Wilhelms-Universität Münster Institut für Geophysik Corrensstraße 24 · 48149 Münster, Germany

Hilmar von Eynatten Geowissenschaftlichen Zentrum der Universität Göttingen Abteilung Sedimentologie/Umweltgeologie Goldschmidtstraße 3 37077 Göttingen, Germany [email protected] Matthias Franz Martin-Luther-Universität Halle-Wittenberg Institut für Geowissenschaften Von-Seckendorff-Platz 3 06120 Halle, Germnay [email protected] Hans-Joachim. Franzke Technische Universität Clausthal Institut für Geologie und Paläontologie Leibnizstraße 10 38678 Clausthal-Zellerfeld, Germany [email protected]

Fabian Hese Christian-Albrechts-Universität zu Kiel Institut für Geowissenschaften, Geophysik Otto-Hahn-Platz 1 · 24118 Kiel, Germany [email protected] Norbert Hoffmann Bundesanstalt für Geowissenschaften und Rohstoffe Dienstbereich Berlin Wilhelmstraße 25-30 · 13593 Berlin, Germany [email protected] Christian Hübscher Universität Hamburg Institut für Geophysik Bundesstraße 55 · 20146 Hamburg, Germany [email protected] Fabian Jähne Friedrich-Schiller-Universität Jena Institut für Geowissenschaften Wöllnitzer Str. 7 · 07749 Jena, Germany [email protected]

Dirk Gajewski Universität Hamburg Institut für Geophysik Bundesstraße 55 · 20146 Hamburg, Germany [email protected]

Hartmut Jödicke † Westfälische Wilhelms-Universität Münster Institut für Geophysik Corrensstraße 24 · 48149 Münster, Germany

Reinhard Gaupp Friedrich-Schiller-Universität Jena Institut für Geowissenschaften Burgweg 11 · 07749 Jena, Germany [email protected]

Marleen Kästner Universität Hannover Institut für Geologie Callinstraße 30 · 30167 Hannover, Germany [email protected]

Marion Geißler Technische Universität Bergakademie Freiberg Institut für Geologie und Paläontologie Bernhard-von-Cotta-Straße 2 09596 Freiberg, Germany [email protected]

Hubert Kiersnowski Polish Geological Institute Rakowiecka 4 00-975 Warszawa, Poland [email protected]

Heijn van Gent RWTH Aachen University Geologie-Endogene Dynamik Lochnerstraße 4-20 · 52056 Aachen, Germany [email protected] Hans-Jürgen Götze Christian-Albrechts-Universität zu Kiel Institut für Geowissenschaften, Geophysik Otto-Hahn-Platz 1 24118 Kiel, Germany [email protected]

Jonas Kley Friedrich-Schiller-Universität Jena Institut für Geowissenschaften Burgweg 11 · 07749 Jena, Germany [email protected] Charlotte M. Krawczyk Geozentrum Hannover Institut für Geowissenschaftliche Gemeinschaftsaufgaben (GGA-Institut) Stilleweg 2 30655 Hannover, Germany [email protected]

Authors Bernhard M. Krooß RWTH Aachen University Lehrstuhl für Geologie, Geochemie und Lagerstätten des Erdöls und der Kohle Lochnerstraße 4-20 · 52056 Aachen, Germany [email protected]

Anna Maria Marotta University of Milan Department of Earth Sciences, Section of Geophysics L. Cicognara 7 · 20129 Milan, Italy [email protected]

Peter A. Kukla RWTH Aachen University Lehrstuhl für Geologie und Paläontologie Wüllnerstraße 2 · 52056 Aachen, Germany [email protected]

Yuriy Maystrenko GeoForschungsZentrum Potsdam · Section 4.3 Telegrafenberg, C424 · 14473 Potsdam, Germany [email protected]

Berit Legler RWE Dea AG Wietze E & P Laboratory Geosciences Industriestraße 2 · 29323 Wietze, Germany [email protected] Rouwen Lehné Johannes Gutenberg-Universität Mainz Institut für Geowissenschaften Becherweg 21 55099 Mainz, Germany [email protected] Christof Lempp Martin-Luther-Universität Halle-Wittenberg Institut für Geowissenschaften, Ingenieurgeologie Von-Seckendorff-Platz 3 06120 Halle, Germany [email protected] Ralf Littke RWTH Aachen University Lehrstuhl für Geologie, Geochemie und Lagerstätten des Erdöls und der Kohle Lochnerstraße 4-20 · 52056 Aachen, Germany [email protected] Tina Lohr GeoForschungsZentrum Potsdam Sektion 3.1 Telegrafenberg, D 225 · 14473 Potsdam, Germany [email protected] Volker Lüders GeoForschungsZentrum Potsdam Sektion 4.3 Telegrafenberg B226 · 14473 Potsdam, Germany [email protected]

Stanislaw Mazur GETECH, Kitson House Elmete Hall Elmete Lane, Leeds LS8 2LJ, UK Markus Mohr RWE Dea AG Überseering 40 · 22297 Hamburg, Germany [email protected] Peter Möller GeoForschungsZentrum Potsdam Telegrafenberg · 14473 Potsdam, Germany Rainer Müller Technische Universität Clausthal Institut für Geologie und Paläontologie Leibnizstraße 10 38678 Clausthal-Zellerfeld, Germany [email protected] Jörg Mutterlose Ruhr-Universität Bochum Institut für Geologie, Mineralogie und Geophysik Universitätsstraße 150 · 44801 Bochum [email protected] Susanne Nelskamp RWTH Aachen University Lehrstuhl für Geologie, Geochemie und Lagerstätten des Erdöls und der Kohle Lochnerstraße 4-20 · 52056 Aachen, Germany [email protected] Georg Nover Rheinische Friedrich-Wilhelms-Universität Bonn Steinmann Institut für Geologie, Mineralogie, Paläontologie Poppelsdorfer Schloß 53115 Bonn, Germany [email protected]

Hans Machel University of Alberta Department of Earth & Atmospheric Sciences 1-26 Earth Sciences Building Edmonton, Alberta, Canada · T6G 2E3 [email protected]

Robert Ondrak GeoForschungsZentrum Potsdam Sektion 4.3 Telegrafenberg, B422 · 14473 Potsdam, Germany [email protected]

Fabien Magri GeoForschungsZentrum Potsdam Section 4.3 Telegrafenberg, C426 · 14473 Potsdam, Germany [email protected]

Birgit Plessen GeoForschungsZentrum Potsdam Sektion 3.3 Telegrafenberg, C327 · 14473 Potsdam, Germany [email protected]

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XX

Authors Rolando di Primio GeoForschungsZentrum Potsdam Section 4.3 Telegrafenberg, B428 · 14473 Potsdam, Germany [email protected] Wolfgang Rabbel Christian-Albrechts-Universität zu Kiel Institut für Geowissenschaften, Geophysik Otto-Hahn-Platz 1 · 24118 Kiel, Germany [email protected] Dirk Radies OMV (Norge) AS Jåttåvågveien 7B 4020 Stavanger, Norway [email protected] Klaus Reicherter RWTH Aachen University Lehr- und Forschungsgebiet Neotektonik und Georisiken Lochnerstraße 4-20 · 52056 Aachen, Germany [email protected] Sabine Rodon RWE Dea AG Überseering 40 · 22297 Hamburg, Germany [email protected] Magdalena Scheck-Wenderoth GeoForschungsZentrum Potsdam · Sektion 4.3 Telegrafenberg, C423 · 14473 Potsdam [email protected]

Judith Sippel GeoForschungsZentrum Potsdam Sektion 4.3 Telegrafenberg, C427 14473 Potsdam, Germany [email protected] Frank Sirocko Johannes Gutenberg-Universität Mainz Institut für Geowissenschaften Becherweg 21 · 55099 Mainz, Germany [email protected] Christopher J. Spiers Utrecht University Department of Earth Sciences Budapestlaan 4 · 3584 CD Utrecht, The Netherlands [email protected] Werner Stackebrandt Landesamt für Bergbau Geologie und Rohstoffe Stahnsdorfer Damm 77 14532 Kleinmachnow, Germany [email protected] Harald Stollhofen RWTH Aachen University Lehrstuhl für Geologie und Paläontologie Wüllnerstraße 2 52056 Aachen, Germany [email protected]

Peter Schikowsky Universität Leipzig Institut für Geophysik und Geologie Talstraße 35 · 04103 Leipzig, Germany [email protected]

David Tanner Geowissenschaftlichen Zentrum der Universität Göttingen Abteilung Strukturgeologie und Geodynamik Goldschmidtstraße 3 37077 Göttingen, Germany [email protected]

Zsolt Schléder Midland Valley Exploration Ltd 144 West George Street Glasgow G2 2HG, United Kingdom [email protected]

Friedrich Theilen † Christian-Albrechts-Universität zu Kiel Institut für Geowissenschaften, Geophysik Otto-Hahn-Platz 1 24118 Kiel, Germany

Sabine Schmidt Christian-Albrechts-Universität zu Kiel Institut für Geowissenschaften, Geophysik Otto-Hahn-Platz 1 · 24118 Kiel, Germany [email protected]

Janos L. Urai RWTH Aachen University Geologie-Endogene Dynamik Lochnerstraße 4-20 · 52056 Aachen, Germany [email protected]

Jörg Schneider Technische Universität Bergakademie Freiberg Institut für Geologie und Paläontologie Bernhard-von-Cotta-Straße 2 09596 Freiberg, Germany [email protected]

Silke Voigt Universität zu Köln Institut für Geologie und Mineralogie Zülpicher Straße 49a 50674 Köln, Germany e-mail: [email protected]

Robert Schöner Friedrich-Schiller-Universität Jena Institut für Geowissenschaften Burgweg 11 · 07749 Jena, Germany [email protected]

Thomas Voigt Friedrich-Schiller-Universität Jena Institut für Geowissenschaften Burgweg 11 · 07749 Jena, Germany [email protected]

Authors John K.Warren Shell Chair in Carbonate Studies (Subsurface Reservoir Characterisation) Sultan Qaboos University PO Box 17 · Al-Khodh-123 Muscat, Sultanate of Oman [email protected] Sven Willert Christian-Albrechts-Universität zu Kiel Institut für Geowissenschaften, Geophysik Otto-Hahn-Platz 1 · 24118 Kiel, Germany [email protected] Jutta Winsemann Universität Hannover Institut für Geologie und Paläontologie Callinstraße 30 30167 Hannover, Germany [email protected]

Tamara Yegorova National Academy of Sciences of Ukraine Institute of Geophysics Palladin av. 32 · 03680 Kiev, Ukraine [email protected] Mi-Kyung Yoon Universität Hamburg Institut für Geophysik Bundesstraße 55 · 20146 Hamburg, Germany [email protected] Henning Zöllner Universität Leipzig Institut für Geophysik und Geologie Talstraße 35 · 04103 Leipzig, Germany [email protected] Christian Zwach StatoilHydro Global Exploration Technology Drammensveien 264 · 0240 Oslo, Norway [email protected]

XXI

Editors

Ralf Littke RWTH Aachen University Prof. Dr. Ralf Littke Institute of Geology and Geochemistry of Petroleum and Coal Lochnerstr. 4-20 · D-52056 Aachen, Germany Phone: +49 241 8095748 Fax: +49 241 80 92152 E-Mail: [email protected]

Ulf Bayer GeoForschungsZentrum Potsdam Telegrafenberg C · 14473 Potsdam, Germany Phone: +49 (0)331 2881340 Fax: +49 (0)331 2881349 E-Mail [email protected]

Ralf Littke received his Diploma degree (M.Sc.) in Geology in 1981 and his Dr. rer. nat. doctoral degree (Ph.D.) in 1985, both from the Ruhr-University Bochum. Afterwards, he worked in the Institute of Petroleum and Organic Geochemistry at the Research Centre Juelich, a major federal research centre in Germany. He taught at the Ruhr-University, Bochum, where he received his Habilitation degree, becoming an Adjunct Professor of Geology in 1993. In 1997, he accepted a Professorship in Geology and Geochemistry of Petroleum and Coal at RWTH Aachen University. His research is focussed on petroleum and gas geology and geochemistry, basin modelling, coal geology and environmental geochemistry. He is coordinator of the German priority research programme ”Dynamics of Sedimentary Basins under varying Stress Regimes (DFG SPP 1135)” and member of the Academy of Science of North Rhine-Westphalia.

Ulf Bayer received his Diploma degree in Geology in 1975 at the Technical University Stuttgart and his Dr. rer. nat. in 1977 at the University of Tübingen. After a time as research assistant he received his Habilitation degree and lectureship there in 1983. In 1981 the Hermann-Credner-Price of the DGG was awarded to him. From 1984 to 1988 he was Heisenberg Fellow of the DFG (German Science Foundation) with terms as guest scientist at the University Leicester and Birmingham (UK), New Brunswick, Princeton and the Research Centre of Schlumberger-Doll at Richfield (USA). From 1988 to 1992 he worked in the Institute of Petroleum and Organic Geochemistry at the Research Centre Jülich and since then at the GeoForschungsZentrum Potsdam. In addition he holds a professorship at the Freie Universität Berlin since 1994. His research focuses on basin analysis and modelling with focus at the integration of geological and geophysical concepts as well as coupled fluid, heat and mass transfer. From 1998-2001 he was a member of the German EUROPROBE steering committee, coordinator of a bundle of DFG-projects related to DEKORP Basin’96 from 1996 to 2001, and is co-coordinator of the German priority research programme ”Dynamics of Sedimentary Basins under varying Stress Regimes (DFG SPP 1135)”.

Freie Universität Berlin Institute for Geological Sciences Malteserstr. 74 -100 · 12249 Berlin, Germany

XXIV Editors

Dirk Gajewski University of Hamburg Prof. Dr. Dirk Gajewski Institute of Geophysics Bundesstr. 55 D-20146 Hamburg, Germany Phone: +49 40 42838 2975 Fax: +49 40 42838 5441 E-Mail: [email protected]

Susanne Nelskamp RWTH Aachen University Dipl. Geow. Susanne Nelskamp Institute of Geology and Geochemistry of Petroleum and Coal Lochnerstr. 4-20 · D-52056 Aachen, Germany Phone: +49 241 80 95779 Fax: +49 241 80 92152 E-Mail: [email protected]

Dirk Gajewski holds the chair of Applied Seismics at the University of Hamburg, Germany where he is employed since 1993. Prior to that he was an assistant professor at the University of Clausthal, Germany and a post doctoral researcher at Stanford University, California, USA and at the Center for Computational Seismology, Lawrence Berkeley Lab, Berkeley, California, USA. He received his Diploma degree (M.Sc.) from the Technical University of Clausthal in 1981 and the Dr. rer. nat. doctoral degree (PhD) from the University of Karlsruhe, Germany in 1987. His research interests are in reflection seismic processing and imaging, seismic anisotropy, and ray methods. He served as associate editor for Geophysical Prospecting, is co-coordinator of the German priority research programme ”Dynamics of Sedimentary Basins under varying Stress Regimes (DFG SPP 1135)” and is a member of the steering committee for the Geophysical Instrument Pool Potsdam (GIPP). Since 2007 he is director of the Wave Inversion Technology (WIT) consortium.

Susanne Nelskamp has studied geosciences at the University of Hannover with an emphasis on Geology. She received her diploma degree (M.Sc.) in 2004 on the topic of benthic foraminifera living in colony with methane reducing bacteria. Up to now she is working on her Dr. rer. nat. (Ph.D) thesis on basin modelling in the Netherlands at the RWTH Aachen University and doing coordination work for the DFG Special Priority Programme 1135.

1

Chapter 1

Characteristics of complex intracontinental sedimentary basins

1

Chapter 1

Characteristics of complex intracontinental sedimentary basins U. Bayer · H.-J. Brink · D. Gajewski · R. Littke

1.1 Introduction The term “basin” is used in different sciences such as physics, chemistry, mathematics, and the geosciences. It can be generally defined as a low, attracting some process coming to rest at the deepest point or the point of minimal energy. A complicated or complex basin then is a potential surface consisting of several minima or basins separated by saddles. A challenging problem is to find out, which minima will be approached (Despa and Berry 2001; Reddy and Pratap 2003; Van Siclen 2007). There is no problem in transferring this rather general definition to geosciences: a basin is a topographic low attracting the accumulation of fluids and sediments by gravity, providing in addition a high preservation potential and, therefore, an excellent archive for reconstructing the past. A basin is classed as complex if several sub-basins exist and interact, e.g., in terms of different hydrological regimes (e.g., Dahan et al. 2004), or with respect to the “migration rules” for oil and gas (IGC 1995), or if along the margins of a rather simple extensional basin sediments are trapped at antithetic faults. These “traps” provide complexity in the form of small sub-basins. Potentials, however, may also change their position and shape through time due to external forces, well known in physics and chemistry where temperature or catalysis may cause essential modifications of the topography of the potential. Similarly in geosciences, eustatic sea-level changes affect the continuity of sedimentation and cause more complex sedimentary patterns (Howell and Flint 1996), and tectonic processes may modify the heat flux as well as the geometry of basins through time (Mazurek et al. 2006) causing complex structures in the sedimentary fill. Finally, an original low may develop into a high. In this case, the basin is inverted and subsequently provides a source of sediments rather than a sink. A basin, whether in chemistry, physics or geosciences, becomes complex either by its geometrical configuration, by involving multiple lows and highs or by temporal changes due to the action of external forces. The latter are typically well defined in the so-named exact natural sciences; however it is more complicated to describe them

in the geosciences. The classification of sedimentary basins, therefore, was and still is an ongoing task. In current text books (e.g., Allen and Allen 2005; Miall 2000) the attempts at a classification appear either as too simplified in terms of one dominating aspect or rather complicated because several aspects and causes are intermixed. In contrast to physics which relies on, for example, “PVT” (Pressure-Volume-Temperature), it is indeed a problem in geosciences to reduce the relevant number of external forces to a well established set. Temperature and gravity drive plate tectonics, causing local tectonics as well as contributing to global climate changes, all feeding back to basin evolution. Furthermore, we cannot reduce basin evolution to the “PVTt-path”, as other factors such as chemical disequilibria also play a role. Here, we provide some facts which are common to complex basins in terms of tectonics, crustal structure, fluid inventory and sediment fill.

1.2 Classifications of basin complexity In order to classify a basin as complex, it must comprise various aspects considered as typical with respect to geophysics, structural geology, sedimentology, geochemistry, etc. Although some of these elements are also present in simple basins and related to relatively simple processes, the overall complexity renders the division of basin evolution into simple elements or basic components difficult. The following sections are an attempt to put these basic components into some logical order.

1.2.1 Tectonic processes – The plate tectonics approach Any approach to classifying basins will depend on the specific scientific background, the aims of “why and how to classify”, and of course the data available: which can come from the sedimentary record, can be based on deep seismic data or can be mainly determined by a geodynamic background. Figure 1.1 illustrates the plate tectonic approach discussed below.

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U. Bayer · H.-J. Brink · D. Gajewski · R. Littke

Figure 1.1. Classic plate tectonic concepts of simple sedimentary basins

As part of the plate tectonic context, the Wilson cycle provides a simple model of the origin, evolution, and disappearance of oceanic basins on Earth. At the beginning, the rift develops due to up-doming of the asthenosphere. Owing to the vertical and horizontal deformation, the crust is stretched and thinned. Subsequently, a larger basin develops and the process will be attenuated due to the load of the accumulating water and sedimentary fill as well as the long term cooling of the asthenosphere. A current example may be found in the East African Graben System, and especially in the Red Sea rift. Not all of the original rifts, however, may develop into oceans; they may get stuck at a certain point and remain simple grabens. If the crust is sufficiently thinned, volcanic events occur which are associated with the crustal thinning: first rheolithes and ignimbrites may occur due to melting of the crust in relation to the increased temperatures; as the stretching continues, mantle derivatives like basalts may take over. This situation is found in several young Graben structures. An example is the Oberrhein Graben (Upper Rhine Valley). The term “Graben” for this structure was introduced by Suess

(1909), thus providing the type locality for a “graben”. However, in detail the Oberrhein Graben is not simple. The graben formation started in the Eocene and during its further evolution several changes in the stress regime occurred causing differences in subsidence and uplift phases, partly confined to certain sections of the Graben area (Illies 1975, 1978; Schumacher 2002; Derer 2003). In addition, pre-existing structures may have affected the local evolution within the Graben area (Schumacher 2002). Another example is the Baikal Rift, which is relatively young (about 30 million years). Currently the Baikal Rift is dominated by pure extension, with mantle flow being horizontal and normal to the rift axis (Gao et al. 2002). However, Petit and Déverchère (2006) claim there was an early stage of formation which was dominated by strike-slip tectonics, creating primitive basins different from the present day ones. They relate the kinematic change from strike-slip to extension with the kinematic reorganisation of Asia due to the growth of the Tibetan plateau. That provides an analogy with the Oberrhein Graben, which also developed in temporal and spatial relation to the formation of the Alps.

Characteristics of complex intracontinental sedimentary basins

In an extended rift zone of several hundred or thousand of kilometers length, the opening process will not be continuous, due to the fact that it takes place on a spherical world. Strike - slip faults develop within the oceanic crust, partially extending into the continental crust causing transtensional and/or transpressional settings wherever a lateral offset occurs. The transtensional setting will cause a subsidence centre due to local extension as may be represented by the Roenne Graben associate with the Tornquist zone, a major lithospheric unconformity at the northern margin of the Central European Basin System. The transpressional system primarily causes local uplift, indirectly providing local minima where sediments can accumulate. As a typical strike slip dominated basin on continental crust we can consider the Jordan rift including the Dead Sea region, or in more general terms the Levante Rift System ranging from the northern Red Sea to southern Anatolia (Mart et al. 2005). The Levante Rift comprises a series of structural basins over more than 1000 km, separated by threshold zones. In general, the Levante rift is characterised by oblique displacement, with dip slip and strike slip at faults. Another such basin is the small but famous Loch Ness, located along the Great Glen fault dividing the North of Scotland along a line from Fort William to Inverness. On a spherical earth, lithosphere has to be consumed elsewhere when an ocean is opening. The subduction of lithosphere and the associated formation of deep sea grabens are well documented along the margins of the Pacific and parts of the Indian Ocean. The local sediment supply will determine whether the sedimentary basin is almost-filled or starved under this compressional setting. Even if filled, the basin will soon become complex. As the sediments are less easily subducted and because of their low specific weight, they are compressed, folded and overthrusted along the resisting plate. In addition, in an oceanic – continental compression setting intra-mountain basins may develop due to the formation of thrusts and napes. The sediment starved Death Valley is an example of this basin type. Further classes of basins associated with compression are back arc and molasse basins located at the rear front of mountain belts. Back arc basins are typically associated with volcanism, crustal melting and thinning while molasse basins are mainly due to flexure of the crust under compression. The “prototype” of molasse basins is the south German and Swiss molasse located at the northern front of the Alps, directly connected to the Alpine compression causing buckling of the continental lithosphere. Its basin evolution was not simple and the Swiss section can be considered complex (Mazurek et al. 2006). Over time, subsidence and uplift phases are recognised, the latter causing erosion of several hundred meters of sediments during post-Cretaceous times. In addition, temperature sensitive data indicate clear temporal changes of the heat flow.

Two classes of basins remain to be mentioned: basins developing under oblique subduction such as those located along the San Andreas Fault System and the so-named sacks represented, e.g., by the Michigan and Illinois basin, and perhaps, the Paris basin. Basins under oblique subduction develop due to a combination of compression and strike slip, combining at least two of the considered principle constituting forces. Concerning the San Andreas Fault, offsets caused the formation of en-echelon arranged pull-apart basins as in the case of a right-stepping offset in Cholame Valley (Shedlock et al. 1990). Material contrasts along the main fault greatly influenced basin evolution (Chavarria et al. 2003). The sacks are still difficult to explain. The areas were or are subsiding, mostly as more or less circular structures without obvious sign of horizontal activity due to compression or extension. Frequently the structures dip rather gently towards the centre and may have lasted relatively long in geological terms as slowly subsiding areas. Several causes have been considered, including (i) underplating from rising asthenospheric material, (ii) alteration of lithospheric material due to the loss of fluids and, therefore, increase of density causing isostatic subsidence, (iii) pure thermal subsidence following a phase of thermallyinduced uplift associated with erosion which provided the sedimentation space during the following contraction and basin formation, or (iv) metamorphism of the lower crust due to a thermal event, followed by volume reduction of the affected rocks and the development of an initial topographic low (Brink 2005a). Due to their often longlasting history and overall low subsidence rates, sacks are greatly affected by facies changes, flooding and erosion due to e.g., global sea level changes or effects of plate movements at great distance. For most basins, there is a lack of the deep seismic experiments which could lead to a better understanding of the processes leading to basin formation. Within this context, complex basins are those characterised by a temporal succession of several of the abovementioned processes such as extension/rifting, strike-slip movements, compression, or inversion within the same region. Complex basins often consist of various sub-basins aligned along a major structure, having somewhat different subsidence and inversion histories. If such basins finally integrate into a larger structure, they may be alternatively termed compound basins. Complications also arise if a stuck rift is later inverted, a characteristic of old structures mostly referred to as Aulacogenes. As the examples mentioned above show, it is difficult to find ideal prototypes for conceptual-derived basin-forming mechanisms. This applies even more to complex basins. Slight changes in the governing stress field or other mechanisms may change the mode of basin formation within

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U. Bayer · H.-J. Brink · D. Gajewski · R. Littke

geologically short times. Clearly, basins that were active over hundreds or millions of years are prone to high complexity in terms of changes in driving processes.

trasts causing stress localisation whereby subsidence and uplift alternated in the sub-regions due to small changes in the primary stress direction.

1.2.2 Crustal association – The strain localization approach

The Pripyat-Dnieper- Donets-Donbas basin is the largest rift in Europe. It originated in the Late Devonian, possibly due to back-arc extension in response to subduction of the Palaeo-Tethys (Lyngsie et al. 2007). It was uplifted in the Permian and in its southernmost part, the Donbas, finally inverted in the Late Cretaceous (Stovba 1996; Stovba and Stephenson 1999). A deep seismic experiment was carried out along a line crossing the Donbas in 1999, consisting of wide-angle refraction and reflection seismics (DOBREflection’99 working group 2002). Both experiments indicate an asymmetric high velocity body below the basin (Fig. 1.2). This may have acted as a ramp during inversion (Maystrenko et al. 2003), leading to a lystric detachment. Together with a major backthrust, a crustal-scale pop-up structure or “horst” developed. However, due to a lack of dating possibilities, it remains uncertain whether the lystric detachment may have originated already during early extension with reactivation during compression. Lyngsie et al. (2007) interpret the reflective lower crust as alternating layers of ultramafic cumulates emplaced in a crust of pyroxene or anorthositic composition. They further assume that the heterogeneous crust evolved during basin extension. This would explain why the rift evolved almost symmetrically rather than showing elements of a Wernicke type of stretching. The later formation of the lystric detachment is most probably due to strain localisation along the more rigid intrusive complex as originally interpreted by Maystrenko et al. (2003).

Basic models of basin formation were developed by considering a homogenous crust, for which mechanical principles are easier to apply than for a heterogeneous layered crust. In reality, however, the lithosphere of the Earth is heterogeneous. This is most obvious in the material contrasts between the oceanic and the continental lithosphere. The oceanic lithosphere, sufficiently far from ridges, is relatively thin but with respect to elasto-mechanic properties it is stiff, whereas the continental lithosphere is at least in parts soft and the deformation follows different rheological laws (e.g., Jaeger and Cook 1971; Meissner 1986). This matter is partially discussed in Cacace et al. (this volume). In addition, the continental lithosphere generally developed over billions of years. Different segments of varying composition were added over time due to plate tectonic processes of rifting and accretion. This is at present observed in the Aegean. The continental lithosphere, therefore, is not just softer than the oceanic one, it is in itself even more heterogeneous due to material contrasts. This applies especially to the uppermost part of the lithosphere, the crust, which mostly was affected repeatedly by compression, extension and strike-slip movements causing additional heterogeneous structures. It is a serious shortcoming that the amount of data decreases the deeper we look in terms of basin formation and basin evolution, or crustal structure. In basin areas there are very few dedicated deep seismic or seismological experiments available. An exception is the DESERT experiment across the Dead Sea Transform (Janssen et al. 2004), which indicates that the seismic basement is offset by 3-5 km and that there is a strong lateral asymmetry in sub-horizontally-layered lower crustal reflectors. Similar structures are observed under the San Andreas Fault. There the differences in crustal structure may be due to the left-lateral movement of about 105 km along the fault and the formation of a series of associated basins. The Oberrhein Graben provides another example where crustal heterogeneities may have affected the origin and further evolution (Schumacher 2002). The deep crustal structure evolved during late Hercynian wrench tectonics (Stephanian, earliest Permian) forming a system of ENE striking troughs and ridges which probably played a role in dividing the Graben area into three sub-provinces (northern, middle, southern) during its evolution. Here, the inherited Palaeozoic structures provided rheological con-

The basins considered so far are relatively simple and illustrate how deeper crustal heterogeneities may cause tectonic regionalisation within the basin area either related to a short time event like the initial extension or later inversion or, sometimes, during its entire development. The relatively young Pannonian Basin may be regarded as another interesting example of complex sedimentary basins. From the viewpoint of structure and tectonic evolution, there are three main aspects which reflect its complexity: 1. The presence of pronounced lateral variations in lithospheric structure and sedimentary basin configuration. The Pannonian Basin as a whole represents an area of pronounced lithospheric weakness with regard to the surrounding Carpathians orogenic front system and foreland (i.e. Bohemian Massif, East European/ Moesian Platform). This marked rheological contrast, directly related to the crustal configuration and thermal properties of the different domains, has induced the high level of strain localisation observed in the Pannonian region. According to Falus et al. (2000) and Seghedi et al. (2004), it may have been affected

Characteristics of complex intracontinental sedimentary basins

Figure 1.2. Crustal structure across the Dnieper-Donets Basin in the Donbas Fold Belt: interpreted reflection seismic above, wide angle velocity model below (modified after Maystrenko et al. (2003), Lyngsie et al. (2007). The development of a large scale lystric fault is related to a high velocity body in the lower crust

in addition by the presence of rising mantle material (mantle plume). 2. Complex polyphase tectonic evolution with repeated activations of pre-existing structures related to the prerift (Late Cretaceous) rheological configuration of the lithosphere and a final inversion phase during the late stage of basin development (Pliocene-Quaternary). The Pannonian Basin evolved under an extensional tectonic setting during Early- to Late-Miocene (about 20-5 Ma). Interestingly, its formation and evolution were marked by a strong coupling between backarc extension of the basin and subduction dynamics across the Carpathians orogenic arc system (e.g., Wortel and Spakman 2000; Cloetingh et al. 2005). This mechanical coupling between the basin, the orogen and its foredeep may explain the anomalous presence of positive inverted structures which are not found in the other Mediterranean back-arc basins. 3. Complex pattern of ongoing tectonic activity. The present-day tectonic setting is the result of large variations, both in time and space, which marked the structural style of the late stage of basin inversion and reflects similar variations in the regional stress and strain fields. High levels of compressional stresses

(up to >100 MPa according to Cloetingh et al. 2005) concentrated within the weak lithosphere of the basin, inducing strain localisation beneath the basin, significant fault activity, increased level of seismicity and differential vertical movements. Considerations concerning inherited heterogeneous crustal structures have not played a major role in basin models for a long time although it is well known from engineering sciences that a small heterogeneity may induce the growth of a fracture and finally lead to failure. In car design, for example, a proper distribution of soft deformable structures and stiff resisting ones are nowadays central for designing a “safe” car in case of an accident. Basins react in similar ways: any contrast in crustal composition will cause strain localisation. Dependent on the general stress field, these heterogeneities determine the locations for deformation in extension, compression or strike slip regimes. The process may be self enforced with respect to changes in the stress field over time, for example if different crustal domains are brought together during larger scale lateral dislocations. There are examples which provide strong evidence that inherited heterogeneous structures influence basin development, causing local

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and/or temporal evolution patterns (see above). However, these examples are still rare due to: (i) a lack of deep crustal studies in most basins, (ii) inadequate modelling techniques, which mostly relied either just on the sedimentary fill or extremely simplified assumptions like homogenous structures, (iii) an absence of communication between the different geosciences involved, namely geophysics, structural geology, geochemistry, and sedimentology. The theme of lithospheric complexity and inherited structures is reconsidered in Part 3 of this book.

1.2.3 Sedimentary systems – The sedimentology approach Complexity in terms of sedimentary fill may be best defined as predictability of the spatio-temporal occurrence of certain facies types and associations. A layer-cake structure of semi-parallel strata would provide the perfectly ordered system. However, this pattern does not really exist in nature and - as was pointed out in Payton (ed. 1977) - an actualistic viewpoint is more realistic. A sedimentary system consists of source areas, transport systems and distal depositional areas. Even if such a complex system were not affected by external forces, the vertical succession would not be uniform as the basin fills with sediments. The succession of layers would indicate “regression” due to the changing topography caused by the process itself – erosion and sediment accumulation. Further complications arise due to the interplay of tectonic basin subsidence and eustatic sea level changes (Vail et al. 1977). In principle, facies distribution is predictable as soon as the relevant processes of sediment distribution in the respective depositional environment are quantitatively understood. After three decades of modern seismic and sequence stratigraphy, this is almost the case for major global sea level changes. Other new methods allow further evaluation of local tectonic processes and an increase in the predictability of their influence on stratigraphy and facies patterns. Nevertheless, we have to realise that we do, at best, understand major depositional processes quantitatively. Therefore, subtle changes in sedimentary successions will remain difficult or impossible to predict. This fact is well known from detailed studies on near-surface aquifers or on petroleum reservoirs which are know to be characterised by subtle changes in mineralogical and petrophysical properties on a meter- and sub-meter scale (Vargas-Guzmán and Al Quassab 2006; Zappa et al. 2006). Complexity may arise in all sedimentary systems, but some are especially prone due to an inherent unpredictability. Fluvial systems which provide stochastic sedimentation and erosion systems in time and space certainly fall into this category. Short time events like exceptional weather situations cause sedimentary patterns which can

hardly be anticipated or reconstructed – neither on short time scales of years nor on a long term geological scale of millions of years. High porosity units of gravel and sand seem to be randomly distributed within silt and clay units. Although the distribution of these generally small scale units cannot be reconstructed, the spatial occurrence and temporal evolution of such a facies complex remains predictable, if major processes and parameters are understood. Complexity becomes a problem of scale, i.e., we have to pose the question of at which time and space level a “facies” is defined. Fluvial, and similarly aeolian, systems provide obviously complex sedimentary patterns. However, even marine systems can be highly complex, e.g., with the occurrence of coarse-grained turbidite and tempestite deposits in paralic, fine-grained successions. Nevertheless, at some distance to continental margins, open ocean deposition is clearly more predictable than continental deposition in terms of their spatial and temporal position with regard to large scale systems. Continental margins on the other hand, tend to be complex, due to active tectonics, abundance of catastrophic mass and turbidity flows, changing deltaic sedimentation, and very complex topography. Complex basin topography (Sinclair 2000) leads to the presence of a set of depositional centers, thus causing spatially variable onlap patterns. An example of the interaction of fluvial and oceanic systems is the present-day Black Sea, in which the sedimentation is complex due to inflow of rivers, basin topography, and currents (Shimkus and Trimonis 1974). The present-day Black Sea sediments are often regarded as an analogue of petroleum source rocks deposited in the geologic past. Especially at sufficiently small scales, the patterns may be further complicated due to diagenetic processes. On a larger scale, salt movements, clay diapirs, and volcanic eruptions act as additional post-depositional sedimentary processes adding to the inherent complexity of sedimentary systems. Generally, inverse density stratification in a sedimentary sequence will cause redistribution of the original layers and thereby modify the topography of the basin floor, leading to small scale heights prone to erosion and nearby rim synclines acting as depocenters. The process frequently becomes self enforced and salt structures which were intitially growing or uprising, may be cannibalised later on by others nearby, showing down-building in later stages. Sedimentation, therefore, may be rather patchy in an environment where local sedimentation and erosion interact at geologically short time scales. A reconstruction of the sedimentary history during an active phase of salt or clay diapir evolution may therefore become a regional puzzle making it impossible to transfer the results from one location to a nearby one. Step-by-step work is required on extremely small scales, using seismic

Characteristics of complex intracontinental sedimentary basins

interpretation, structural reconstructions and modelling, as well as interpreting all types of borehole and core data in order to unravel the history of sedimentation around just one salt diapir. Summarising the above sections we have identified two major classes of basin complexity in sedimentary systems: 1. Complexity due to the basic processes which trigger erosion and sedimentation such as (i) occasional floods spilling sediments into depressions, (ii) channels meandering and changing the landscape (in fluvial, lacustrine, deltaic or submarine systems), and consequently permanently modify the basin topography by generating new deposition and potential erosion centers. 2. Complexity due to post-depositional processes such as diagenesis, salt- and clay diapirism. On a temporal scale deposition centers may be eroded and generated. Thus, the sedimentary system becomes even less predictable.

1.2.4 Fluid and mineral inventory – The diagenetic and/or petroleum system approach Sedimentation is followed by internal alterations due to increasing pressure and temperature during burial. Basins evolve as geochemical reactors, thus varying the rock and fluid properties of the sedimentary infill in the course of time. This function as a “geochemical reactor” leads to the formation of groundwater, oil, natural gas, and coal as well as many other fluid and mineral resources. Sedimentary basins can, therefore, be regarded as the greatest resource of mankind. With respect to the “geo-reactor” concept of sedimentary basins, temperature evolution over time is of utmost importance, because temperature is the major driving force for processes affecting basin fill. Accordingly, the age of a basin is an important attribute. For example, a basin that developed early in the Earth’s history may have been affected by several geodynamic processes during its own history, resulting in a high grade of complexity with respect to generated and accumulated fluids. The “geochemical reactor” especially modifies the physical properties of rocks, in part on a rather small scale. In addition, the chemical characteristics of pore-water may change due to chemical reaction with the host rock, generating new fluids. This means the fluid inventory can also be used to characterise sedimentary basins. Simple, shallow basins are characterised by having water as pore fluid, because no hydrocarbons were generated. Deeper, older, and more complex basins usually show a more differentiated

fluid inventory. Original pore waters are partly replaced by salty brines or by inflowing meteoric waters. In addition, different types of hydrocarbons may exist, including black oil, condensate and natural gas. Furthermore, generation of inorganic gases is favored, carbon dioxide, molecular nitrogen and hydrogen sulfide being among the most prominent. Complex basins can contain all these fluid types in varying quantities. Their occurrence greatly depends on fluid transport pathways (aquifers; migration routes) which are zones of high permeability. Such zones are commonly separated by low permeability zones (aquitards, seals, migration barriers). If such barriers are very effective, different fluid compartments may exist within a single sedimentary basin. The integrity/stability of such compartments over long (geologic) periods of time is a matter of ongoing research. There is, however, no basin classification scheme based solely on the “geochemical reactor”. Indeed, the reactor is controlled by the entire basin history and basin research is driven mainly by the search for oil and gas within the fill of a sedimentary basin. The attempts to classify basins in terms of their petroleum content, therefore, are complex in themselves. The distribution of hydrocarbons in ensembles of fields of different sizes, plus the requirement for the better prediction of such systems as function of basin types, links the geological framework to commercial considerations. Important contributions in that sense are basin classification methods that are related to the oil and gas distributions. These distributions are seen as (statistical) projections of (geological) basin characteristics and can therefore be used for classification purposes. An essential aspect thereby is the prediction of field-size distribution within a given basin, an attempt going back to Klemme (1984). He analysed 65 of the world’s basins representing 85% of the world’s discovered and produced petroleum and concentrated on the larger fields in each basin (Fig. 1.3). Significant variations in the percent-rank of the five largest fields appear to be fundamentally related to the basin’s architectural form or morphology (a major element in basin classification). Second order variations within the morphology/size relation of the basin’s first five fields’ size-distribution appear to be dependent on basin fill and the fact that the classification system results in many basins falling into more than one category (hybridism). Four first order categories were applied: (i) craton interior basins, (ii) continental multi-cycle basins, (iii) continental rifted basins, and (iv) delta basins. The five largest fields of craton interior basins and delta basins are generally very similar and contain only a small portion of the basins’ hydrocarbon reserves. These two observations mean that incremental differences between the ranked fields of the petroleum system are small and that the system itself with its very similar members can be

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characterised as relatively simple. In the other two categories, continental multi-cycle basins and continental rifted basins, the concentration of hydrocarbon reserves within the five largest fields varies according to basin size. These variations are significant and probably represent higher complexity of the hydrocarbon system. Thus, the complexity of hydrocarbon systems can be used as a rough classification tool for basins since in many respects the petroleum system also represents the overall geological complexity. Basins with more than one hydrocarbon system can be subdivided at least laterally into different units with their individual complexity. Based on Exxon’s 600 identifiable sedimentary basins worldwide, Kingston et al. (1983) proposed a highly complex classification scheme. The objective of this exercise was to establish valid comparisons of tectonic style and, therefore, of basins: their sedimentary fills, oil plays, and potentials. The basic unit in this classification is the cycle, which consists of the sediments deposited during one tectonic episode. Some basins have only one sedimentary or tectonic cycle. These are then called simple basins. Most basins, however, contain more than one tectonic/sedimentary cycle, and are called polyhistory basins.

Using the petroleum system of a basin as a characterisation tool, Demaison and Huizinga (1991) proposed another method. Their genetic classification of petroleum systems is based on a simple working nomenclature that consists of combining qualifiers from each of the following three categories: (i) charge factor (supercharged, normally charged, or undercharged), (ii) migration drainage style (vertically drained or laterally drained), and (iii) entrapment style (high impedance or low impedance). In this sense the North Sea Basin, at least its British and Norwegian parts, are high impedance, vertically drained, supercharged basins, reflecting the tendency of a rift basin to be vertically drained due to petroleum transfer along faults and fracture systems (Fig. 1.4). The necessary elements of a petroleum system are source rocks, migration pathways, reservoir rocks, seals and traps as well as overburden rocks as necessary elements. Relative timing of formation of these elements as well as of petroleum generation and migration are of utmost importance for accumulations to form and to be preserved. A chart showing these elements is shown as Figure 1.5, where geologic time is plotted along the horizontal axis and the petroleum system elements along the

Figure 1.3. Comparison of size of different basin types with percentage of petroleum found in the five largest fields (after Klemme 1984)

Characteristics of complex intracontinental sedimentary basins

Figure 1.4. Example of a vertically drained, high impedance petroleum system in a rift basin (example from the Central Graben, North Sea, United Kingdom, after Thomas et al. 1985) Petroleum is expected to move preferentially in a vertical direction along fault and fracture systems into reservoirs. Petroleum is generated in the source rock, but only below the dashed line indicating the top of the oil window. Buoyancy-driven transport is upwards, until traps are reached (after Demaison and Huizinga 1991)

Figure 1.5. Modified Petroleum System Chart for the Carboniferous-Rotliegend gas system of the Central European Basin system (after Gautier 2003; see Ulmishek and Klemme 1990 for further explanation)

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vertical axis (Ulmishek and Klemme 1990). Exploration strategies should be developed in and restricted to sedimentary basins in which a complete petroleum system is expected to exist. Basins that are complex with respect to fluid systems are often characterised by multiple source rocks, e.g., deep gas source rocks and more shallow oil source rocks, and/or multiple trap types, e.g., anticlines, fault related traps, stratigraphic traps, salt related traps. Furthermore, some of these basins have reached temperatures sufficient for thermochemical sulphate reduction resulting in the generation of hydrogen sulfide, or have access to molecular nitrogen or carbon dioxide sources, e.g., from deep crustal or even mantle sources. Clearly, the older and tectonically more complex a sedimentary basin is, the more likely is a variable inventory of fluids. Fluids, especially hydrocarbons, and diagenetic processes provide a seemingly random process which may turn any simple basin type into a complex basin with regard to exploration for resources. Any attempt towards a straightforward prediction would require a detailed knowledge of the entire history, which cannot be achieved, and can at best only be approximated.

1.3 Summary We have defined basin complexity as the interaction of several of the above discussed aspects or approaches. Following Webster’s New World Dictionary complex is also defined as a group of inter-related structures or processes, or psychologically as an integration of ideas related to a particular object. Both aspects are also important in geosciences: 1. The basin may be complicated in terms of repeated tectonic events which are not easily subdivided: it may have evolved over a heterogeneous crust, causing strain localisation in specific areas due to tectonic forces; it may comprise a complex stratigraphic fill due to repeated localised erosion and sedimentation or primarily complex sedimentation; finally diagenet-

ic processes may be inhomogeneous and fluids may further complicate local structures. This definition of complexity as the interaction of several of the above discussed aspects or approaches was mentioned right at the beginning of the chapter. 2. The second aspect, the psychological or intellectual integration of different techniques and different geosciences, has not been dealt with. This aspect becomes particularly manifest in the “oil plays” where an attempt is made to integrate theoretical models and data or, more precisely, concepts from geophysics, tectonics, sedimentology, and inorganic as well as organic geochemistry. This task in itself is complex. Scale is another aspect to be considered. The plate tectonic approach represents the largest scale comprising thousands of kilometres, and the diagenetical approach represents the smallest scale in this classification. These scales, however, may interact in various ways with each other. As an example, a fluid migration path along a fracture system caused by strain localisation of a plate tectonic process may form an oil play at a salt tectonically-related trap or a buried fluvial system. This example indicates that complexity and scale are related. This is hardly surprising seeing as we have defined complexity as the interaction of more than one of the above aspects according to scale. Scale is also linked to age since diagentical processes require certain pressure and temperature conditions which are usually not present when the basin is initialised. If we define complexity by the ability to predict a system it is vital in which way we can describe the processes steering the formation of a basin. Whereas large scale processes like graben formation are usually described by deterministic models based on simplified assumptions, small scale features like fluvial systems, turbidite flows or some diagenetic processes may be only described by a stochastic approach. These statistics are usually “statistics of small numbers” resulting in large error bars due to the small data base. It is well known that the transfer of a petroleum system from one location to another even within the same basin does not work at all. This unpredictability of the system can also be considered as another measure of complexity.

Characteristics of complex intracontinental sedimentary basins

Box 1 What is complex? -> Integrating groups of inter-related structures or ideas 1. Plate tectonics Basics: Complex:

a. Strike-Slip associated (transtensional/transpressional) b. Rift Systems (Wilson cycle, failed rifts) c. Sags (e.g. under plating, metamorphic processes) d. Compressional settings of various types Several types superimposed in time

2. Crustal association Basics: a. Rheology (ocean/continent) b. Lateral variations in the crust or lithosphere (e.g. thickness) c. Strain localisation due to lateral variations in material properties Complex: No simple relationship between stress and strain, strain localisation 3. Sedimentary fill Basics: Complex:

a. Complex onlap-offlap patterns b. Multiple local erosion and sedimentation c. Salt and mud diapirs disturbing the sequence and complicating sedimentation Complexity may be inherent or the sedimentary record hides true basin evolution

4. Diagenesis and fluid sytems Basics: a. Multiple types of fluids b. Random distribution in field size c. Chemical sub-provinces d. Convective fluid movements e. Small scale diagenetic alterations, fluctuating physical properties Complex: Predictability of resources and their quality Complexity: Several aspects apply to one basin or a compound basin Of course, complex means also: a. That different scales interact, from 100rds of km down to µm b. That temporal successions of various processes overprint original signals c. That chaotic non deterministic events may be involved during deposition (e.g., turbidites) as well as inside the basin (e.g., fluid convection)

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Chapter 2

The Central European Basin System – an Overview

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Chapter 2

The Central European Basin System – an Overview Y. Maystrenko · U. Bayer · H.-J. Brink · R. Littke

2.1 Introduction The geological knowledge of the Central European Basin system (CEBS, Fig. 2.1) is founded on more than 1,000 years of industrially exploring and exploiting mineral resources therein. At its southern rim within the Harz Mountains mining of Palaeozoic rocks for silver and other metals is well reported for at least one millennium, terminating the prospering Viking silver trade from Southeast Europe which is seen as one of the causes for the decline of the Viking power. By reconnoitering the mining industry of the Harz for the Duke of Sachsen-Weimar in the late 18th century, Johann Wolfgang von Goethe made some important contributions to a rational understanding of the

geology of the area. In the “Osnabrücker Bergland” of the inverted Lower Saxony Basin industrial ore mining began in the 12th century (Rose 1984) with the quarrying of silver bearing galena of the Zechstein. The mining of Carboniferous coals started in the Ruhr district in the 14th century and, the mining of Wealden coals of the Lower Saxony Basin in the 15th century. Locally the Wealden coals were used for the evaporation process of Zechstein salt out of saline waters. Zechstein salt itself had already been industrially exploited by the city of Lüneburg north of Hannover since the 10th century. Since no coals for burning were available nearby for the evaporation of saline waters, wood was used and large diluvial areas were deforested in the “Lüneburger Heide” during medieval times. Quarries

Figure 2.1. Topography of central Europe with outlines of the Early Permian (Rotliegend) basins (modified after Lokhorst et al. 1998), showing Permian basin depocentres in relation to present-day relief. Early Permian basins: NPB – Northern Permian Basin, PT – Polish Trough and SPB – Southern Permian Basin

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of Cretaceous and Middle Triassic limestones (e.g., Helgoland) and of Zechstein dolomites at few specific saltplug related sites within the North German Lowlands proved the presence of rocks within basin, which were already known from its fringes. In the course of the Industrial Revolution of the 19th century, the exploration and ex-

and these have been assessed in terms of their hydrocarbon prospectivity. Exploration wells and reflection-seismic data are the main information sources. Indirect data on deep horizons that were not penetrated by wells are provided by gravimetric, magnetic, magneto-

Figure 2.2. Oil field Wietze, 1910 (courtesy of the German Oil Museum – Deutsches Erdölmuseum Wietze)

ploitation of coals and lignite, ores and salt increased and deeper drilling for research was carried out. In 1858/59 a well was drilled in the search for lignite in the village of Wietze north of Hannover (see Fig. 2.2), encountering an oil bearing layer. In this area oil seeps had already been detected in the 17th century, and the oil was used as cart grease in those times (Ehrenwerth 1998). In the following decades up until present day, national and international companies have successfully explored and exploited significant amounts of hydrocarbons from the CEBS. The history of two of those companies is well documented by Fahrion/BEB (1987) and Mayer-Gürr (1994). Exploration for oil and gas, including numerous wells and geophysical data acquired since the early 20th century, and its production have provided a major contribution to our knowledge of the subsurface of the CEBS down to a depth of about 5 – 8 km. The up to 10 km thick sequence comprises mainly Quaternary to Rotliegend sediments

telluric, refraction-seismic and particularly important, deep seismic reflection lines (Dohr et al. 1989; Krawczyk et al. this volume). Until the beginning of the 1990s, commercial 2D seismic lines with recording lengths up to 14 seconds were acquired on a large scale (several thousand km). As far as acquisition and processing were concerned, the lines were predominantly adjusted to address questions specifically related to hydrocarbon exploration. The maximum distances between the geophone locations and shotpoints were between 1,800 m and 4,800 m for the deep seismic lines. With the increased use of 3D seismics with shorter recording times, usually exclusively acquired above gas fields, no more new 2D seismic lines were recorded. A long period of oil price decline in the 1990s further contributed to the fact that virtually no more deep seismic lines were acquired. Nevertheless, scientific application of refraction/reflection seismic profiling since the early 1990s has made essential contributions towards the understanding of the basin structure.

The Central European Basin System – an Overview

The European Geotraverse was a first attempt to reveal the deep structure across the transition zone between the Caledonian/Variscian domain in NW Germany and Baltica (Blundell et al. 1992; Blundell 1999; EUGENO-S Working Group 1988). The deep structure of Baltica and the Sorgenfrei-Tornquist Zone was investigated by the BABEL project (BABEL Working Group 1993). In 1993 seismic data were acquired by the MONA LISA project in the south-eastern North Sea in order to study the structural relation between the Caledonian collision zone and Late Palaeozoic to Mesozoic basins (MONA LISA Working Group 1997a,b). An integrated reflection and refraction line (DEKORP Basin’96) across the Northeast German Basin yielded new insights concerning the structure and evolution of the Southern Permian Basin (DEKORP-BASIN Research Group 1999). The deep seismic sounding experiments POLONAISE’97 and CELEBRATION 2000 aimed to clarify the deep structure of the Tornquist-Teisseyre Zone (TTZ) and the Polish Trough (PT; Guterch et al. 1999, 2003). Finally, the temporary seismic station array (TOR) was designed to study the lithosphere–asthenosphere system across the northwestern part of the TransEuropean Suture Zone by teleseismic tomography during 1996-97 (Arlitt et al. 1999; Gregersen et al. 2002). The Central European Basin System (CEBS) covers the area from the North Sea to Poland and from Norway to the German midlands (Fig. 2.1). Although having its origin in the latest Carboniferous/ earliest Permian, the basin area is still characterised by lowlands. The central part of the CEBS corresponds to the Dutch, North German and Danish Lowlands which are characterised by low topography, 25 m above sea level on average. The adjacent areas of the North and Baltic Seas also have relatively shallow bathymetry. Notable changes in relief occur along the southern margin of the CEBS, where present-day relief is elevated north of the Harz Mountains and the Sudetes as well as in Poland (Holy Cross Mountains). Scandinavia bounds the CEBS area in the north. The present-day topography mainly reflects the latest evolutionary stage of a 300 Ma (million years) long history since the Late Carboniferous-Early Permian. On the other hand, the geological history of the area has been recorded since the Ordovician with several tectonic cycles within the Caledonian (Pharaoh 1999; Krawczyk et al. 2008) and Variscan (McCann et al. 2006; Kröner et al. 2008) orogenies. At the transition from the Carboniferous to the Permian three sub-basins developed, the Northern and Southern Permian basins as well as the Polish Trough (Fig. 2.1). During subsequent evolution, the CEBS was significantly deformed by both external and internal tectonic forces evolving as a geochemical reactor that reworked the rock and fluid properties of the sedimentary infill over time.

The geodynamic processes, the properties of the underlying crust, the areal size, the volume and composition of the sedimentary load, the type of internal structure and the distribution of generated hydrocarbon and inert fluids define the characteristics of this basin system. All these characteristics are time dependent, defining the age of the CEBS as an important attribute as well. Clearly, the CEBS represents one of the most important records of the Earth’s history during the last 300 million years.

2.2 Crustal association The CEBS covers an assemblage of crustal domains of different consolidation age in the transition area from Precambrian Baltic Shield and East European Craton to Caledonian–Variscan Europe (Fig. 2.3a). The present-day crustal association of the CEBS is the result of long lasting processes of terrane amalgamation on the Precambrian Baltic and East European Craton during the Precambrian and Palaeozoic. The major fault zone of the Trans-European Suture Zone is the Tornquist Zone which includes two components, the Tornquist-Teisseyre (TTZ) and the Sorgenfrei–Tornquist zone (STZ). The Tornquist-Teisseyre Zone separates Precambrian East European crust from Palaeozoic crust of Central Europe (Berthelsen 1992), while the Sorgenfrei–Tornquist Zone was identified as a major lithosphere boundary (Gregersen et al. 2002, 2005; Shomali et al. 2006). The area between the Elbe-Odra Line (EOL) and the Sorgenfrei–Tornquist Zone exhibits sharp and steep steps at the Moho level and the lithosphere/asthenosphere transition as schematically illustrated in figure 2.3b (Shomali et al. 2006; Thybo 2001). In order to understand the present-day structure of the lithosphere, the early tectonic history of the area has to be considered. During the Late Ordovician-earliest Silurian, East Avalonia accreted to Baltica (Fig. 2.4a; Cocks et al. 1997) causing the formation of the Caledonian Deformation Front (CDF in Fig. 2.3a) which is documented by seismic lines as a zone of thrusting (e.g., Tanner and Meissner 1996; Abramovitz et al. 1998; Krawczyk et al. 2002). This collision between Avalonia and Baltica marked the last stage of Laurussia suturing (Fig. 4b; Ziegler 1990). Since the 1990s, the Thor Suture (or Caledonian Deformation Front at the shallow level) has usually been considered as the collision suture between Baltica and East Avalonia (Pharaoh et al. 1997) while Cocks et al. (1997) preferred the Elbe-Odra Line as the major contact zone. Furthermore, a model of a wedge of Baltic crust continuing from the Ringkøbing-Fyn High to the Elbe-Odra Line was proposed by DEKORP-BASIN Research Group et al. (1999) and favoured by Bayer et al. (2002). Aichroth et al. (1992), Bayer et al. (2002) and Yegorova et al. (2007b)

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Figure 2.3. (a) Major crustal structures of the transition zone between Baltica and the East European Craton and Caledonian-Variscan Europe (modified after Bayer et al. 2002 and Pharaoh 1999). Structural elements: CDF – Caledonian Deformation front; DSHFZ – Dowsing–South Hewett Fault Zone; EEC – East European Craton; EFS – Elbe Fault System; EOL – Elbe-Odra Line; LU – Lysogory Unit; MM – Malopolska Massif; RS – Rheic Suture; STZ – Sorgenfrei–Tornquist Zone; TS –Thor Suture; TTZ – Tornquist-Teisseyre Zone; VF – Variscan Front. (b) Schematic cross-section showing crust-mantle and lithosphere-astenosphere configurations across the major structural domains of the CEBS (simplified after Shomali et al. 2002)

have shown that the area between the Thor Suture in the north and the Elbe-Odra Line in the south is characterised by high velocities (6.8 – 7.0 km/s) and high densities in the lower crust resembling the similar structured Baltic lower crust. In contrast southward, from the Elbe-Odra Line to the Variscides, a low-velocity zone in the lower crust has been recognised (Thybo 2001; Scheck et al. 2002b). Similar results concerning the Elbe-Odra Line as the southernmost boundary of Baltica were derived from the TOR experiment which indicates a regional thinning of the lithosphere from the Baltic Shield towards the ElbeOdra Line (Gregersen et al. 2002, 2005; Shomali et al. 2006). Therefore, it was concluded that the wedge of high

velocity lower crust provides a relict of the former passive margin of Baltica which was overthrust by Avalonia where the Caledonian deformation front is the northern shallow termination of the Thor suture. The Rheic Suture (RS) separates East Avalonia and other terranes from the Variscan domain which accreted in Late Palaeozoic time. This suture is associated with the closure of the Rheic Ocean and is in part well imaged by seismic reflection data as a crustal shear zone (Pharaoh 1999). The collision related processes within the Rheic Ocean caused the Variscan Orogeny which culminated during Early to Middle Carboniferous when the supercon-

The Central European Basin System – an Overview

Figure 2.4. Global plate reconstructions from the Early Ordovician (a) through the Late Silurian (b) to the Early Carboniferous (c) (modified after Scotese 2004)

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tinent Pangea was forming (Fig. 2.4c; Stampfli and Borel 2002). The extent of Variscan thrust sheets is expressed by the Variscan Front (VF) which is characterised by a sequence of arcs (Fig. 2.3a). In contrast to Proterozoic and Caledonian Europe, the Variscan internides are characterised by a low-density crust and upper mantle (Yegorova et al. 2007b), implying another compositional and/or thermal regime. The differences between Precambrian and Palaeozoic Europe are highlighted by the magnetic field (Fig. 2.5) which is characterised by chaotic and relatively shortwave anomalies over the East European Craton and Baltica, while long-wave anomalies dominate over the Palaeozoic consolidated European crust. The complex accretion history of the area led to an assemblage of varying crustal rheologies below the CEBS area. This crustal differentiation then played an important role during the post-Variscan evolution of the CEBS causing strain and stress localisation in areas, where major changes in the structure and composition of the crust and in part the upper mantle occur.

2.3 Permian Basin formation and subsequent subsidence In the Early Permian the supercontinent Pangea was assembled (Fig. 2.6a) and the intracontinental Permian basins originated during Latest Carboniferous-Early Permian times when the CEBS was affected by extensive igneous activity and faulting (Gast 1988; Plein 1990; Ziegler 1990; Dadlez et al. 1995; Bachmann and Hoffmann 1997; Bayer et al. 1999; Abramovitz and Thybo 1999). The basin system is superimposed on the Westphalian foreland basin which developed in front of the Variscides and follows their curved front (Baldschuhn et al. 2001) with a general SWNE strike (Fig. 2.7). The CEBS in contrary strikes NW-SE. This difference in basin strike implies essential changes of the stress regime during the Late Carboniferous-Early Permian. Furthermore, the WNW-ESE elongated Southern Permian Basin matches the orientation of the Elbe-Odra Line and the Elbe Fault System (EFS in Fig. 2.7) and is complicated by the NW-SE-striking Polish segment which is parallel to the Sorgenfrei–Tornquist Zone.

Figure 2.5. Total magnetic field of the CEBS, upward continued to 3 km altitude (from the database by Wybraniec et al. 1998) with major crustal boundaries (after Bayer et al. 2002 and Pharaoh 1999) and outlines of the Early Permian (Rotliegend) basins (modified after Lokhorst et al. 1998). Structural elements: TS – Thor Suture, for names of other crustal boundaries see Fig. 2.3

The Central European Basin System – an Overview

Figure 2.6. Global plate reconstructions from the Early Permian (a) to the Late Cretaceous (d) (modified after Scotese 2004)

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Figure 2.7. Major strikes of the Westphalian basin (light yellow arrow) and the Southern Permian Basin and the Polish Trough (black arrows), shown on the Westphalian (green) and the Early Permian (Rotliegend) (grey) depocentres (compiled after Baldschuhn et al. 2001; Bayer et al. 2002; Lokhorst et al. 1998 and Pharaoh 1999). For names of the major crustal boundaries, see Fig. 2.3

According to Ziegler (1990), a major wrench fault was active along the Elbe Fault System in Central Europe during the Late Carboniferous. Geological evidence for Late Carboniferous dextral transtensional movements were reported by Mattern (1996) along the Elbe Fault System. In addition, dextral horizontal movements along the Elbe Fault System are inferred from a pull-apart model for the Late Carboniferous-Early Permian rifting phase of the North German Basin considering the distribution of initial deposits (Bachmann and Grosse 1989; Scheck and Bayer 1999). Furthermore, latest Carboniferous to earliest Permian transtensional faulting occurred along the Tornquist Fan due to dextral strike-slip movement at the Tornquist Zone (Thybo 1997). The Tornquist Fan is defined as a triangular structure between the Sorgenfrei-Tornquist zone and the Ringkøbing-Fyn High. Scheck et al. (2002b) then proposed that the Tornquist Fan and the Elbe Fault System acted as border faults during latest Carboniferous rifting, giving rise to the simultaneous formation of the Southern Permian Basin (North German Basin and Southern North Sea), the Danish Basin, and the Polish Trough in the earliest Permian. The known lithology of the Early Permian (Rotliegend) is characterised by the presence of thick volcanic

rhyolites and ignimbrite series (Benek et al. 1996; Breitkreuz et al. this volume) at its base, indicating that the crust was partially melted (e.g., Gast 1988; Plein 1990; Benek et al. 1996; Bachmann and Hoffmann 1997). Occasionally basaltic material was also encountered at some locations. The total volume percentage of these basalts, however, is very small (about 2%, Benek et al. 1996). Apparently, the Late Carboniferous-Early Permian tectonic event caused metamorphic processes in the lower crust, resulting in increased rock densities and subsequent decrease of rock volume (Brink 2005a,b). According to Brink (2005a), the volume reduction processes resulting from metamorphism can explain roughly 30% of the subsidence, and the remaining 70% are related to the sedimentary load. Furthermore, the thermal anomaly in Late Carboniferous-Early Permian times is estimated to have lasted about 25 Ma (~300-275 Ma). However, commonly it is still understood by the majority of geoscientists that thermal relaxation of the lithosphere and sedimentary loading dominated subsidence during the Permian (e.g., Ziegler 1990; Scheck and Bayer 1999; Van Wees et al. 2000).

The Central European Basin System – an Overview

2.4 Subsequent formation of sub-basins During the Triassic, the CEBS was affected by several pulses of extensional tectonics which led to the formation of superimposed sub-basins (Fig. 2.8). At the regional scale, the Triassic corresponds to a period of global plate reorganisation, marking the beginning break-up of Pangea (Fig. 2.6b; Ziegler 1990). The most intense subsidence occurred in the sub-basins or grabens surrounding the Ringkøbing-Fyn High, in the Glückstadt Graben (GG, in Fig. 2.8; Baldschuhn et al. 2001; Maystrenko et al. 2006), the Horn Graben (HG; Clausen and Pedersen 1999; Baldschuhn et al. 2001) and the Fjerritslev Trough (FT; Britze and Japsen 1991). Other centres of Triassic subsidence are located within the Sole Pit Basin, the Central Graben (CG), the Rheinsberg Trough (RT) and the Polish Trough (PT; Van Hoorn 1987; Ziegler 1990; Dadlez 2003; Evans et al. 2003; Scheck et al. 2003a,b). The Polish Trough and the Fjerritslev Trough followed the direction of the Tornquist–Teisseyre Zone and the Sorgenfrei–Tornquist Zone, respectivly, and therefore, these basins inherited Permian basin orientation controlled by crustal heterogeneities along the Tornquist Zone. On the other hand, the NE-SW-striking Glückstadt Graben, the Horn Graben and

the Rheinsberg Trough did not follow the general trend of the Southern Permian Basin which was more or less NW-SE (Fig. 2.8). In addition, the Central Graben cuts the Northern and Southern Permian basins from NNW to SSE. Therefore, differentiation of the Permian basins into several Triassic sub-basins together with changes in orientation of the Triassic depocenters marks a first important complication of the CEBS during post-Permian time (Cacace et al. this volume). After the Triassic, the Pangaea break-up (Fig. 2.6c) was accompanied by Late Jurassic-Early Cretaceous rifting in several areas of the CEBS leading to the Lower Saxony Basin (LSB; Betz et al. 1987; Jordan and Kockel 1991), the Central Graben (Oakman and Partington 1998; Erratt et al. 1999; Nielsen 2002; Moeller and Rasmussen 2003), the West Netherland Basin (Van Wijhe 1987; Duin et al. 2006) and the Polish Trough (PT; Dadlez 2003; Lamarche and Scheck-Wenderoth 2005). In the latest Cretaceous Africa–Arabia and Eurasia collided (Fig. 2.6d), causing the formation of the Alpine orogen. The collision-induced compressional stresses affected the whole of Europe, with some basins undergoing compression and inversion (Kley et al. this volume; Voigt et al. this volume). Again, zones of strong hetere-

Figure 2.8. Main post-Permian sub-basin superimposed on the Early Permian (Rotliegend) basins (grey) (compiled after Bayer et al. 2002; Kockel 1988; Lokhorst et al. 1998; Maystrenko et al. 2005a,b; Pharaoh 1999; Scheck et al. 2003a,b; Vejbaek and Britze 1994; Ziegler 1990). Yellow area: Tertiary subsidence centre in the North Sea Basins: CG – Central Graben; FT – Fjerritslev Trough; GG – Glückstadt Graben; HG – Horn Graben; LSB – Lower Saxony Basin; PT – Polish Trough. For names of the major crustal boundaries, see Fig. 2.3

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ogeneities in the crust caused strain localisation during compression along the Elbe Fault System, the Teisseyre-Tornquist Zone and the Sorgenfrei-Tornquist Zone (Ziegler 1990; Scheck et al. 2002b; Otto 2003; Scheck-Wenderoth and Lamarche 2005; Mazur et al. 2005; Krzywiec 2006b). Localisation of the strongest compressional deformations led to the formation of deep reaching reverse faults and flexures, causing significant erosion of sediments at the southern margin of the CEBS and along the Tornquist Zone during the Late Cretaceous-Palaeogene. Therefore, the basin structure was modified in terms of inversion, marking the final stage of basin development in particular areas of the CEBS, this being vertical uplift and subsequent erosion.

The complex structure of the CEBS is highlighted by the gravity map in figure 2.9. The gravity signals within the CEBS correlate well with Triassic and Jurassic subbasins and inversion features, reflecting density variations at the upper crustal level due to local sediment accumulation and/or erosion. The inverted structures appear as a NW–SE-striking chain of positive anomalies extending along the Elbe Fault System. Positive anomalies also mark the Central Graben of the North Sea and the Rheinsberg Trough of the North German Basin. On the other hand, the Fjerritslev Trough (FT, for location see Fig. 2.8), the Horn and Glückstadt Graben are characterised by gravity lows. These lows can be partially explained by huge salt structures within the thick sedimentary cover.

The Late Cenozoic is characterised by the formation of a broad depocentre within the North Sea (Fig. 2.7) in contrast to the marginal parts of the CEBS (Sclater and Christie 1980; Ziegler 1990; Nielsen 2002) with subsequent subsidence throughout the Cenozoic. Especially since the Miocene, more than 2,000 m of sediments have accumulated in the central North Sea area. There is some indication that subsidence continues locally. However, it is difficult to distinguish between tectonically-induced and salt-induced subsidence.

2.5 Sedimentary history The sedimentation history of the area was affected by a variety of factors acting at different scales, e.g., plate tectonic movements through various climate zones, global sea level changes, large scale tectonics such as the breakup of Pangea, more localised tectonics such as the already mentioned strain localisation during inversion and very local features related to salt tectonics. The general outline

Figure 2.9. Gravity anomaly map (Bouguer anomalies onshore and Free-Air anomalies offshore) for the CEBS (from the database by Wybraniec et al. 1998). For names of the structural elements, see Fig. 2.3

The Central European Basin System – an Overview

Figure 2.10. Generalised stratigraphic-lithologic columns for the (a) northeastern, (b) central, and (c) southwestern part of the CEBS (after Resak et al. 2007; Rodon and Littke 2005; Senglaub et al. 2006). Signatures according to Shell Standard Legend (1995), stratigraphy according to Gradstein et al. (2004)

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given here, therefore, can only be a rough sketch. Typical stratigraphic-lithologic columns for the eastern, central and western onshore part are presented in Fig. 2.10 (Littke et al. this volume). In the Late Carboniferous (Westphalian), the area of the CEBS was located more or less at the equator and inhabited by tropical forests and abundant swamps and raised bogs leading to the formation of coal seams. This situation changed when Pangea moved further northward into the arid zone during the Permian. Red, oxidised clastics now dominated. Under these conditions sedimentary cycles consisting of clastics and salts developed, reflecting minor changes in aridity (Permian Rotliegend Formation; Bachmann et al. this volume). The Rotliegend clastics are the most important gas reservoir rocks of the basin and have been strongly modified by diagenetic processes (Schöner et al. this volume). The situation changed when a connection to the open sea developed between Norway and Greenland in the Late Permian (Zechstein) and the depression was flooded. Still under arid conditions, several typical evaporite cycles were deposited, starting with carbonates through gypsum to halite (Warren, this volume). Flooding obviously occurred repeatedly (Ziegler 1990; Plein 1990; George and Berry 1997; Evans et al. 2003). Salt accumulation was remarkable, reaching more than 2,000 m in the basin centre (Maystrenko et al. 2005a,b). During the Triassic, the climate was still predominantly arid, but the number of humid events increased over time, indicated by fossil soils in the Buntsandstein, thin coal seams in the lowermost Keuper (Lettenkeuper), and especially by river systems which evolved in the Late Triassic (Stollhofen et al. this volume). In addition, a sea way opened temporally in the Muschelkalk connecting the Basin from Germany through Poland to the Tethys, flooding the basin area which in its marginal parts now extended to southern Germany. However, full marine conditions were established at best for very short times as documented by salt deposits and the evolution of endemic faunas like the Ceratides. Furthermore, the Triassic extensional event in the CEBS triggered Permian salt movements and caused the formation of deep rim synclines within the basin area, bringing about local subsidence and complicated internal patterns of Late Triassic sedimentation (Kukla et al. this volume). During the Jurassic the basin area passed 30 °N and encountered increasingly humid conditions. In addition wide areas of Europe were now flooded by shallow seas providing a direct connection of the basin area with the Tethys and fully marine conditions. During two intervals, the Early and the Late Jurassic, black shales were deposit-

ed providing well preserved fossils, but more importantly excellent source rocks for petroleum (di Primio et al. this volume). In contrast, during the Middle and Late Jurassic carbonate deposition became important, including thick oolite sequences associated partly with small coral reefs and micritic carbonates. These periods of carbonate sedimentation were interrupted by periods of clastic sedimentation. However, sedimentation rates remained low to moderate (Bachmann et al. this volume). These facies changes indicate either a high variability of climatic conditions or of clastic supply from the hinterland, perhaps related to an unsteady northward movement of the continent. At the end of the Jurassic, the sea generally retreated, with coral reefs appearing in south Germany and remains of fossil forests in southern England (Portlandian), followed by the Wealden, extending into the Early Cretaceous and causing a widespread unconformity probably related to a major drop in sea level. In addition, the Jurassic rifting phase in the CEBS reinforced the declining Triassic salt movements and caused the formation of sedimentation centres within the marginal areas of the Triassic depocentres, while JurassicEarly Cretaceous subsidence was minor in the areas of most intensive Triassic sedimentation. Furthermore, Late Jurassic to Early Cretaceous erosion occurred within the former Triassic depositional centres, such as the Glückstadt and Horn Grabens. The general subsidence pattern changed, complicating the entire basin system. During Jurassic and Early Cretaceous times the CEBS was under the influence of both warm and cool waters: ammonite fauna indicates two major faunal provinces, the south German one with relations to the Tethys and the British one which sometimes is termed “boreal” (Stollhofen et al. this volume). During this period, the entire basin system was inverted: areas of formerly high subsidence received little sediment or were uplifted, whereas marginal areas with low thicknesses of Permian to Jurassic sediments subsided relatively fast. In contrast these marginal troughs such as the Lower Saxony Basin were strongly uplifted during the Late Cretaceous. In more central parts such as the Pompeckj Block, renewed, sedimentation set in with deposition of chalky limestones produced by calcareous nanofossils. They mark a culmination of the Cretaceous transgression when the central part of the CEBS was covered by seawater. During the Cenozoic, continental to deltaic clastics prevailed within the CEBS which was gradually moving to the north as a part of western Eurasia until it reached the present-day position. The present-day surface was mainly formed by Quaternary glaciations (Gibbard 1988; Sirokko et al. this volume). Thereby an interesting situa-

The Central European Basin System – an Overview

tion arises because the river system followed the front of the ice sheets with a general east-west to SE-NW direction during and right after the glaciations period (“Urstromtäler”, ice-marginal valley). However, within the last 10,000 years the river system adapted once more to the deeply located crustal discontinuities, crossing the ice-marginal valleys frequently more or less south-north. This pattern indicatesl ongoing movements localised at these inherited structures as already noted by Woldstedt (1961, see also Sirokko et al. this volume). Throughout time, the sedimentary cover of the CEBS has been strongly affected by salt tectonics since the Triassic (Fig. 2.11a) providing a large natural laboratory concerning the effects of salt tectonics in space and time. For example, the cross-section through the North German Basin (Fig. 2.11b) displays salt-induced deformations of the overburden which are observed above the approxi-

mately flat base of the Permian salt. It shows that the salt cover was deformed due to salt movements and that the Permian salt layer decoupled deformation of the overburden from the strata below. This means that Permian salt acted as an internal independent factor of basin evolution and contributed significantly to its complexity. Salt as a rather ductile material starts to flow easily under differential pressure (Urai et al. this volume), building underground “landscapes” larger in size than any mountain chain on the Earth. Once initialised, salt flow may evolve into an autonomous process. However, salt may also just rest over millions of years within a basin under continuous sedimentation until it is disturbed by an additional tectonic event. Concerning the present-day distribution of salt structures, two regions are distinguished (Scheck-Wenderoth et al. this volume; Fig. 2.11a): (i) in the central part of the

Figure 2.11. (a) Distribution of the salt structures (blue contours) within the CEBS. Limit of the Zechstein Basin is shown in grey (modified after Lokhorst et al. 1998). (b) Cross-section along the DEKORP Basin’96 reflection seismic line (for location, see Fig. 2.11a) showing the salt-related deformations within the sedimentary cover across the Northeast German Basin (simplified after Mazur et al. 2005)

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Figure 2.12. Distribution of methane and nitrogen in main Permian (Rotliegend) reservoirs within the western and eastern parts of the CEBS (compiled after Bayer et al. 2002; Lokhorst et al.1998 and Pharaoh 1999). For names of the major crustal boundaries, see Fig. 2.3

basin system more or less N-S directed salt walls dominate, while (ii) at the present-day margins, especially the southern one, the salt walls are almost NW-SE orientated. Concerning the primary initiation of salt movement, the N-S direction is related to Triassic extension, while the NW-SE direction reflects localised inversion during the latest Cretaceous. During the Triassic, the first stage of widespread salt diapirism in the CEBS was triggered by regional extension along transtensional lineaments (Brink et al. 1992). Once the salt structures reached the critical size, buoyancy forces supported their continued growth until the Jurassic when extension-induced regional stresses once more affected the CEBS. During the late Early Cretaceous-early Late Cretaceous, less salt movement occurred. Salt domes further evolved during the Cenozoic, reaching the surface at some places in the CEBS. In summary it can be stated that tectonically-induced salt movements had an important impact on sedimentation and the subsequent deformation of Mesozoic and Cenozoic strata within the CEBS. Furthermore, the huge salt structures complicate the fluid and thermal regimes of the entire basin system due to the strong difference in physical parameters (such as almost zero porosity, very low permeability, high thermal conductivity etc.) compared to other sedimentary rocks. In consequence the thermal

field is essentially disturbed within the CEBS, causing for example thermo-haline driven fluid movements or local modification of the maturation of organic matter (Littke et al. this volume; Magri et al. this volume).

2.6 Fluids within the CEBS Concerning the complex evolution of the CEBS it is certainly no surprise that the fluid composition shows an enormous degree of complexity. This is most evident for petroleum fluids which consist of oils and gases of variable density and composition. For example, the major gas reservoir sandstones in the Permian Rotliegend formation contain methane and molecular nitrogen in vastly different percentages with clear regional trends (Krooss et al. this volume, di Primio et al. this volume). This fluid composition is related to the presence and maximum depth of burial of the Carboniferous gas source rocks which in turn depends on crustal structure (Fig. 2.12). A more detailed view is presented in figure 2.13, which shows the distribution of Westphalian source rocks compared to methane concentrations in Carboniferous and Rotliegend gas reservoirs. The presence of a thick coal-bearing Westphalian sequence seems to be a prerequisite for the occurrence of high-quality gas in large parts of the study area.

The Central European Basin System – an Overview

With respect to liquid petroleum in Mesozoic reservoirs, the Toarcian Posidonia Shale is regarded as the principal source rock in the onshore part of the CEBS, except for the eastern (Polish) part, where Zechstein carbonates are important source rocks. There are three major problems with respect to liquid oil generation from the Posidonia Shale: (i) the vast erosion of the Early Jurassic, including the Posidonia Shale, (ii) the facies change from highly oil prone in the west to inert in the east (Fig. 2.14), and (iii) the low maturity of the source rock in parts of the basin (Gaupp et al. this volume). Charge, accumulation, and entrapment reflect the distribution and re-distribution of fluids within the sedimentary basin. Over the course of time, migrating petroleum fluids, which are lighter than water, will be transported upwards and partly escape via seepages towards the surface of the Earth. This transport is greatly controlled by the permeability distribution of the system, its history and the current status. On a worldwide scale Miller (1992) estimated the loss of oil due to seepage and introduced it

into a global system of oil generation, loss, half-life and the world crude oil reserves. Conventional oil in reservoirs has a well-defined half-life of about 29 Ma, meaning that after this time period 50% of the world’s oil reserves would vanish, if they are not replaced by newly generated ones. For the Zechstein gas fields within the German part of the Southern Permian Basin, Brink (2002a) estimated a half-life of about 60 million years based on their leakage along faults through the generally sealing Zechstein evaporites into overlying Triassic Buntsandstein traps. Furthermore, the hydrocarbon generation potential of the km-thick Carboniferous source rock sequence was enormous. Compared to the generated methane, little gas has been trapped (with the exception of the Groningen area; Littke et al. 1995) indicating that by far the greatest part of the gas has been lost from the sedimentary system of the CEBS. This also indicates that the Zechstein salt seal was anything but perfect over the long history of the basin. We return to Klemme’s approach (Bayer et al. this volume) that field size distributions (FSD) can be used for

Fig. 2.13. Distribution of methane contents in main Permian (Rotliegend) and Carboniferous reservoirs and distribution of Westphalian gas source rocks within the CEBS (compiled after Bayer et al. 2002; Lokhorst et al. 1998 and Pharaoh 1999). For names of the major crustal boundaries, see Fig. 2.3

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Fig. 2.14. Toarcian Posidonia Shale organic facies and distribution within the CEBS and location of oil fields (compiled after Bayer et al. 2002; Lokhorst et al. 1998 and Pharaoh 1999). For names of the major crustal boundaries, see Fig. 2.3

the classification of sedimentary basins. Normally, they display log-normal behaviour and represent multiplicative systems. The number of parameters, whose multiplication leads to a field quantity, is represented by the steepness of the log-normal graph (linear approximation of the Gaussian integral). It can be shown by Monte Carlo inversion that the more independent parameters involved, the steeper the graph of a log-normal distribution (Megill 1977, 1979; Brink 2000). The size of a discovered oil or gas field can be determined multiplicatively by its geometric dimensions like area and reservoir thickness, by its porosity and water saturation as well as by other factors. In addition, the probable size of an oil or gas prospect can be estimated by assuming a range of numbers for these parameters and choosing e.g., the mean of a set of independent “Monte Carlo” runs as an estimate of the most likely case. Using an estimate of the number of relevant parameters derived via the rough and simple “Monte Carlo inversion” (Brink 2000), comparisons of the field size distributions of Zechstein and Rotliegend gasfields of the CEBS indicate

different geological complexities and, in addition, different commercial discovery probabilities as well. Depending on the absolute field sizes, the higher the complexity, the smaller the discovery rate will be. Excluding the Salzwedel gas field the normalised gas field distributions in the Permian reservoirs of the Northern German Basin show a very complex Rotliegend system with a large number of independent parameters and a less complex Zechstein system. Within the Rotliegend, the British hydrocarbon system (PGS 1996) can be defined to be more simple than the German one (Brink 2000). A comparison of the field size distribution of the Permian fields in Poland (Zechstein and Rotliegend reservoirs together are part of one system (Bandlowa 1998)) with the German Zechstein and Rotliegend indicates an intermediate complexity that lies in between German distributions (Fig 2.15). The Dutch system of Rotliegend fields is even simpler than the British one. Extracting the necessary information from Van de Weerd (2004) and the Ministerie van Economische Zaken (2002) an estimated FSD for the Rotliegend fields in The Netherlands can be

The Central European Basin System – an Overview

Figure 2.15. Field size distributions of Rotliegend and Zechstein gas fields of the Southern Permian Basin

Figure 2.16. Field size distributions of some selected oil systems for various oil regions

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Y. Maystrenko · U. Bayer · H.-J. Brink · R. Littke

derived. In this case the giant Groningen field has to be excluded first from the field size distribution, as it represents a basin wide anomaly. The remaining largest 50 Rotliegend fields have a spread of a few tens of billion m³, leading to a relatively small increment between neighbouring field sizes. The differences between the single regionally divided systems of the Southern Permian Basin are shaped considerably by the different reservoir and structural parameter distributions as well as by the variations of generation and migration of hydrocarbons due to the area dependent complex burial and thermal histories. The spread of differences may define a further classification method for basins. Published data about the size or final total production of oilfields within the CEBS region exists e.g., for the Lower Saxony Basin (Binot et al. 1993) and the Ekofisk region in the Central North Sea Graben (Cornford 1994). Data on a very interesting distribution for the Niger delta in Nigeria are also available (Thomas 1995). As expected by following Klemme’s (1984) observation, the Nigerian delta basin is characterised by a low dip in its log-normal field size distribution, while the continental rift basin of the North Sea around Ekofisk has a really high log-normal dip and the Lower Saxony Basin (Betz et al. 1987) as part of the CEBS has an intermediate log-normal dip (Fig. 2.16).

2.7. The CEBS – prototype of a complex sedimentary basin The CEBS comprises all the attributes necessary to be considered a prototype of a complex basin system: (i) even during its initial formation three connected sub-basins developed in a not easily understood arrangement, (ii) during its further thermal subsidence phase additional small scale basins were attached and/or superimposed complicating the tectonic story, (iii) a rather differentiated crust and mantle structure was inherited from older processes, providing the basis for strong rheological contrasts within the basin area, (iv) changes in the general stress field and lateral variation in stress field above the salt, (v) sedimentation was not just affected by general sea-level and climatic changes but was complicated by localised tectonic movements through the subsidence history causing both some erosion and in other areas sediment accumulation at the same time and hampering straight forward detailed correlation between subareas, (vi) salt tectonics reacted rather sensitively to any tectonic event and developed its own dynamics, thus complicating the internal structure on a local scale, and (vii) the fluid inventory from water to hydrocarbons is in many parts not predictable or at best only very crudely. Although extensive exploration for hydrocarbons has been carried out in the CEBS over more than 100 years and scientific experiments performed for almost two decades, some parts of the basin history leading to the complex inventory remain a mystery.

3

Chapter 3

Strain and temperature in space and time

3

Chapter 3.1

Driving mechanisms for basin formation and evolution M. Cacace · U. Bayer · A.M. Marotta · C. Lempp

3.1.1 Driving mechanisms for basin evolution A basin may be called “complex” if its temporal evolution was complicated by several tectonic or thermal events such as stretching and inversion or if it is composed of a variety of regionally distributed sub-basins with differing histories. Aulacogens or typically inverted “failed rifts” are examples of the first type. The second type is frequently encountered in intraplate settings located above continental crust. This may be because the continental crust is rather inhomogeneous as it results from long-term accretion of terrains, manifested by different thicknesses and materials. A problem then is our generally very poor knowledge of the deeper crust especially in basin areas with thick sedimentary cover. Evidence of deep structures can only be gathered by seismic methods, well established in mountain ranges (COCORP, DEKORP), but mostly missing in oil or gas prawn basins where exploration focuses on the sedimentary fill. The early physically motivated basin models developed in the 1970th and 1980th, therefore, relied on extreme simplifications concerning the lithospheric structure, temperature distribution with depth, and tectonic scenarios, and can essentially be divided into two concepts: 1. Basins in extensional settings comprising simple rifts and transtensional regimes represented by the kine­ matic models of McKenzie and Wernicke, relying on perfect plasticity of the lithosphere after stretching. 2. Basins in compressional settings like along subduction zones, foreland and back arc basins represented by flexure or buckling of an idealised elastic beam or plate. Although these basic models may well account for some simple young basins, they can at best account for a time shot in a complex intracratonic basin suffering a variety of processes in time and space. Since then, and parallel to the development of computer power, more and more numerical models have been developed, involving complex rheologies and boundary conditions to be solved as non-linear temporal process.

Although in a physical sense they help to understand why a complex basin reacted in a certain way over time, they do not allow the reproduction of data in the way required, e.g., for oil exploration. Actually this was already a problem with the physically motivated kinematic models, and a second line of basin modelling was developed after the 1970s which relied on the sedimentary fill together with chemical control parameters. Here we will rely on physically motivated models and we will use the Central European Basin as an example where most of the concepts presented have been applied to a certain temporal event or a specific region. This will allow us to elucidate the limits of the still rather simple models as well as the problems encountered in complex sedimentary basins. The Central European Basin System is certainly an outstanding example of this type comprising all mentioned spatial and temporal criteria.

3.1.2 Kinematic models for basin formation 3.1.2.1 Purely thermal models Approximately 70% of the surface of the continental crust is covered by more than 2 km of sediments (Sclater and Celerier 1987). The major accumulations occur in sedimentary basins, locally reaching thicknesses of more than 10 km. Sedimentary basins are regions of prolonged subsidence of the Earth’s surface. The origin and characteristics of the driving mechanisms within the mechanical rigid lithosphere have been long debated. One class of models used to explain vertical movements in continental regions closely resembled the thermal model which has been successfully used for oceanic lithosphere (e.g., Vogt and Ostenso 1967). Following this approach, the subsidence of continental shelves could in principle be related to thermal contraction beneath the crust. This conclusion reflects the concept that the tectonic subsidence of continental lithosphere decreases exponentially as a function of time with a time constant very close to that typical of a mid-ocean ridge (Sleep 1971; Steckler and

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Watts 1978). Sleep (1971) proposed a major thermal perturbation as the driving mechanism for subsidence. The thermal anomaly heats the entire lithosphere, thus causing uplifting of the crust by thermal expansion. Subsequent removal of the upper crustal layers by erosion together with the resultant cooling produce subsidence below the original surface level. Although this model accounts for the time history of subsidence rather well, the explanation is inconsistent with the large sediment accumulations frequently observed. When the temperature of the lithosphere increases, the surface is first elevated and then starts to subside to its original position due to the cooling of the lithosphere. Without erosion, this process will not lead to the formation of a sediment-filled basin. The model causes a “space problem”: given the similarity in density between sediments and upper crustal layers, subsidence should match the amount of material eroded. However, there is no evidence for such a great amount of erosion in deep sedimentary basins. Another problem concerns the mechanism required to explain the heating of the continental lithosphere. An attempt to explain this “heating problem” was suggested by Artemjev and Artyushkov (1971). In order to describe the anomalously high heat flow in the Baikal Lake, they proposed an upwelling of hot upper mantle material during the extension phase with a consequent thinning of the crust as the driving mechanisms of the observed high heat flow. Alternatively, Haxby et al. (1976) suggested the heat

flow anomaly was caused by a possible intrusion of hot mantle diapirs into the lower part of the lithosphere which do not produce major deformation at the surface level.

3.1.2.2 McKenzie’s kinematic model In order to avoid both the space and the heating problem, McKenzie (1978a) developed a quantitative model of passive mechanical extension of both crustal and subcrustal lithosphere, i.e., uniform stretching. He considered instantaneous and uniform extension of the lithosphere and the crust with passive upwelling of hot asthenosphere to maintain isostatic equilibrium. Consequently, mechanical stretching is considered responsible for both heating of the lithosphere and subsidence of the basin area. The assumptions related to this model can be summarised as follows:

· · ·

The lithosphere overlies a partially molten asthenosphere. The densities of the crust, lithosphere and the asthenosphere decrease linearly with temperature, whereby the mantle-lithosphere and the asthenosphere present the same thermal properties. The density of the crust, however, is lower than the densities of the asthenosphere and the mantle-lithosphere. The geotherm is considered to be a linear function Figure 3.1.1. Thermal evolution as a function of time as predicted by McKenzie’s (1978a) stretching model. A: Steady state geotherm before stretching. B: Steady state geotherm after stretching. C: Final steady state geo-therm (red thick curve). After instantaneous increase in heat flow, the temperature decreases exponentially with time. The total temperature is made up of a steady and of an unsteady components. Red dotted curves indicate the transient temperature as a function of time

Driving mechanisms for basin formation and evolution

Figure 3.1.2. Principal features of the McKenzie’s (1978a) subsidence model. See box 3.1.1 for the mathematical formulation. A: Initial conditions. A thermally equilibrated continental lithosphere of total thickness tL consisting of a crust of thickness tc and a lithospheric mantle of thickness tm=(tc – tL) overlies a partially molten asthenosphere. B: Uniform instantaneous stretching (β). At the time t = 0, uniform instantaneous mechanical extension of the lithosphere by a factor β occurs causing vertical thinning of both the crust (thickness = tc/ β) and the mantle-lithosphere (thickness = tm/β =(tc – tL)/ β ). Since the temperature of the material remains unchanged during the extension, isostatic compensation causes upwelling of hot asthenosphere. The resulting gradual decaying of the thermal perturbation produces an initial instantaneous subsidence, Si. C: Post-rift evolution. The cooling of the lithosphere following rifting causes a second phase of relative slower thermal subsidence, ST(t)

· · ·

from a fixed temperature at the base of the lithosphere (Ta ~1300 °C) to the surface temperature (T0). Radiogenic internal heat production within the crust is neglected. Initial mechanical stretching, β=original thickness/final thickness, is considered instantaneous and volume is preserved, i.e., homogeneous within the lithosphere = uniform stretching. Lateral temperature gradients are much smaller than vertical gradients, therefore the heat excess is gradually removed only by vertical heat conduction and the effects of lateral heat flow are neglected.

Under these assumptions, the model predicts the temporal thermal evolution (Fig. 3.1.1) as well as the total amount of subsidence (Fig. 3.1.2) which consists of two components (Box 3.1.1):

1. Rapid stretching of the lithosphere produces thinning of both the crust and the mantle-lithosphere and upwelling of hot asthenosphere, perturbing the original thermal structure (Case B of Fig. 3.1.1). The temperature gradient increases of an amount equals to the stretching factor β because crust and mantle are thinned while the temperatures at the surface and at the transition to the asthenosphere remain constant. The thinned lithosphere is replaced by less dense athenosphere causing isostatic readjustment and instantaneous rapid initial subsidence, Si (Case B of Fig. 3.1.2). The predicted initial subsidence depends on the initial thickness of the crust, tc, and is independent of the amount of stretching β. According to McKenzie (1978a), uplift occurs for pre-stretched initial crustal thickness less than 20 km whereas subsidence is predicted for crustal thickness initially greater than 20 km.

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BOX 3.1.1 Mathematical formulation of McKenzie’s (1978a) subsidence model A.Initial condition tL = total initial (un-rifted) lithospheric thickness; tc = initial (un-rifted) crustal thickness; tm = (tL – tc) = initial (un-rifted) mantle-lithosphere thickness; ρc = average density of the crust; ρm = averaged density of the mantle-lithosphere;

B.Instantaneous mechanical stretching (β) Initial (fault-controlled) subsidence, Si:

Si> 0 ⇒ subsidence; Si< 0 ⇒ uplift; β = stretching factor; ρs =average bulk density of sediments or water or air filling the rifted basin; αV=volumetric thermal expansion coefficient (for both the crust and the mantle);

C. Post-rift evolution Thermal temporal evolution:

thermal constant of the lithosphere; z = vertical spatial direction (depth); t = time; k = thermal diffusivity; Thermal subsidence, ST (t):

Driving mechanisms for basin formation and evolution

2. The disturbed temperature gradient causes cooling whereby asthenosphere is transformed into mantle. While the crust remains thinned, the entire lithosphere approaches its original thickness (Case C of Fig. 3.1.2). This cooling process is nonlinear with depth (Case C of Fig. 3.1.1). The relaxation of lithospheric isotherms to their pre-stretching position is associated with a second thermal subsidence phase ST(t) whose subsidence rate decreases exponentially with time. The thermal subsidence depends only on the amount of stretching β. Since the initial and final thermal states are known, the temperature distribution as well as the surface heat flow can be calculated as a function of time. Once the thermal structure is known, the subsidence history of the extended region is easily estimated. The quantitative aspects of the extensional concept were first applied by McKenzie (1978b) to the Aegean Sea. LePichon and Sibuet (1981) applied the stretching model of McKenzie to the formation of passive continental margins, focusing on the Armorican and Galicia continental margins of the northeast Atlantic. They found out that the simple stretching model was able to explain the relationship between initial subsidence, thermal subsidence and continental crustal thinning, thus being a good first approximation of the actual physical process of formation of the margin. Moreover, the simplicity of such an approach makes it possible to reconstruct the edge of the continent prior to break-up and to account for the dynamics of the transition from continental stretching to oceanic accretion. Several other groups have applied these concepts to other basins and shelves with generally successful results (e.g., Royden and Keen 1980; Wood and Barton 1983; Brunet and LePichon 1982).

Figure 3.1.3. Effect of lateral heat conduction on subsidence history. Lateral heat conduction (red curve) provides a faster subsidence than the one predicted by McKenzie’s (1978a) simple stretching model (black curve)

3.1.2.3 Limitations of the McKenzie’s model and corresponding implementations The simple one-dimensional uniform extension model developed by McKenzie provides a quantitative explanation of the general aspects of continental basins evolution. However, many authors have demonstrated discrepancies occurring between the crustal extension, β, initial subsidence and thermal subsidence as predicted by the McKenzie’s model with regard to geological observations. Subsidence analysis of well data from Nova Scotia by Royden and Keen (1980) showed that McKenzie’s model predictions may lead to unrealistic overestimates of the initial subsidence. Additionally, crustal extension is commonly much less than the model predicts. Time-dependent models for lithospheric deformation have demonstrated that a purely mechanical source of extension may not be able to generate the heat flow and uplift history as actually observed in a large number of rift zones. The assumption of instantaneous stretching of the lithosphere followed by thermal subsidence during the re-equilibration stage of the lithosphere is an attractive assumption since it provides a simple initial condition for the thermal calculations, thus simplifying the analysis of the post-rift thermal subsidence of the basin. However, such an approximation is reasonable for modelling a slab of infinite horizontal extension undergoing a stretching event for less than 20 Ma. Prolonged periods of rifting may cause significant amounts of lateral heat loss, thus enhancing the amount of subsidence during the stretching phase. The longer the rifting event, the more important lateral heat loss should become. Additionally, Cochran (1983) suggested that supplementary heat loss away from the basin centre may facilitate increased early basin subsidence and parallel thermal uplift of the basin flanks. To account for lateral heat conduction effects, Steckler (1981) expanded the one-dimensional vertical heat flow model to two dimensions by also including conduction in one horizontal direction (Fig. 3.1.3). Cochran (1983), applying this implementation, developed an analytical technique to investigate the effects of an extended rifting event for the simple case of a sediment free basin, i.e., water filled. The obtained results showed several deviations from the McKenzie’s approach and the finite rifting case in terms of subsidence history, distribution of sediments and thermal history for extension times as short as 5 Ma. The main effect of considering a finite-duration extension event is a heat loss during rifting which leads to an increase in the syn-rift subsidence at the expense of the post-rift, resulting in a more rapid subsidence in the early history of the basin evolution followed by a slower thermal subsidence than is predicted by the McKenzie’s

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Figure 3.1.4. Effect of finite rifting rates on thermal (post-rift) subsidence. Blue curve: end-member subsidence curve for the case of instan-taneous (δt6=0) uniform stretching as developed by McKenzie (1978a). Red curves: subsidence curves for different finite rates of rifting. As a consequence of considering a finite duration of extension, an increase in the syn-rift (initial) subsidence occurs at the expense of the post-rift (thermal) subsidence. This feature becomes more important the longer the rifting event is considered and it is reflected by a progressively flattening of the post-rift subsidence curves. Purple line: end-member subsidence curve for the case of infinitely slow rifting rate (δt1=∞). In this situation the mantle-lithosphere will remain cool and almost no thermal subsidence will occur after the rifting has ceased

model (Fig. 3.1.4). A further consequence of considering lateral heat flow is an increase in total heat flow and subsequent thermal expansion which causes significant uplift across the basin flanks. However, when applied to a sediments-filled basin, the model of Cochran (1983) gives unrealistically large ratios of syn- to post-rift sediments. In summary, McKenzie’s model provides an attractive tool to quantify one-dimensional subsidence using some simple numbers like the stretching factor even if only well data are available. However, it remains a first order approximation with severe restrictions, even for simple basins. It should be used very circumspectly in studies involving complex basins which suffered several phases of subsidence and perhaps inversion as well. It may be easily adapted to approach such a chain of events. The model is limited by the strictly one-dimensional approach which neglects the lateral variations observed in the real world as well as by the simplified assumptions regarding the temperature distribution in the lithosphere and the rheology of the earth.

3.1.2.4 Non-uniform stretching models: discontinuous and continuous stretching with depth The simple geometrically uniform, depth-independent stretching model does not take into account the layered

rheological stratification of the lithosphere. Since rheological properties vary with temperature and pressure, it is more realistic to expect the lithosphere to extend in a non-uniformly manner with depth. The distribution of lithospheric extension with depth may be discontinuous or continuous (Fig. 3.1.5). In the first case, the upper and lower lithosphere are decoupled at a certain depth that may or may not correspond to the crust-mantle boundary. The portion of the lithosphere above this horizon extends by a certain factor δ, which is usually less than the amount the underlying lower lithosphere extends, stretching factor=β. The upper and lower lithospheric extensions are considered independent but uniform throughout their respective layer thickness. Clearly, when δ=β uniform extension occurs. With respect to the uniform stretching model, discontinuous stretching raises the lithosphere-asthenosphere boundary to a shallower level. In this way it solves the “heat problem” since thinning the sub-crustal lithosphere more than the upper crust results in increased heat input during extension. At the same time, discontinuous stretching models require a minor amount of crustal thinning for the same total subsidence. As in the uniform case, the total subsidence is made up of two contributions, an initial fault-controlled subsidence and a subsequent thermal subsidence. The initial subsidence depends on the initial crustal thickness, tc, the decoupling depth and the relative magnitudes of δ and β, whereas the thermal subsidence is hardly affected. Considering the situation of instantaneous stretching, i.e., no heat loss during the rifting process, the thermal subsidence reflects the amount of sub-crustal thinning. This provides a simple way to estimate the stretching factors of the crust and mantle-lithosphere directly from the amount of syn- and post-rift subsidence. Royden and Keen (1980) applied such a model to the Nova Scotian and Labrador continental margins. In their formulation the decoupling horizon was located at the base of the crust and instantaneous extension was assumed. A general agreement was found between the theoretical subsidence predicted by their model and the subsidence calculated from deep well data. Moreover, the results explained the uplift typically experienced by many margins during early rifting phases and show that the thermal subsidence can account for the long-term tectonic subsidence observed. A more refined model allowing for multiple rifting phases as well as differential extension of the crust and the mantle-lithosphere was proposed by Van Wees et al. (2000) to explain the Late Permian-Early Triassic evolution of the Southern Permian Basin, the southern part of the Central European Basin system (CEBS) (Fig. 3.1.6). The results indicate:

Driving mechanisms for basin formation and evolution

Figure 3.1.5. Differences between uniform (A), discontinuous with depth (B), and continuous with depth (C) stretching models. A Uniform stretching model: the crust (initial thickness tc) and the mantle-lithosphere (initial thickness tm = (tL – tc) with tL = total initial thickness of the lithosphere) are stretched by an identical amount (β). B Discontinuous stretching with depth: the crust (initial thickness tc) is stretched by a different amount (δ) with regard to the mantle-lithosphere (initial thickness tm = (tL – tc) which extends by a factor β>δ. This difference in the amount of crustal and sub-crustal extension requires a decoupling between the two layers. Both crustal and sub-crustal extensions are considered independent but uniform throughout the respective layers. C Continuous stretching with depth: in the crust the stretching is the same of the above described situation, while in the sub-crustal layer the stretching is a continuous function of depth

· · · ·

The basin wide Late Permian–Triassic subsidence required an active tectonic mechanism, although there is almost no evidence of synchronous active faulting. A relatively important component of Late Permian and Triassic tectonic subsidence can be explained by thermal relaxation of a major lithospheric thermomechanical attenuation which occurred during Early Permian and by a consequent delayed infilling of the topographic depressions which developed during the Late Permian. The Early Permian stretching was associated with thermal-induced thinning of the mantle rather than mechanical driven crustal extension, consistent with the Stephanian-Autunian phase of wrench tectonics and magmatism. Thinning of the crust below the Southern Permian Basin may be partly attributed to its mechanical ex-

tension and partly to magmatic destabilisation of the crust-mantle boundary followed by reactivation of the lower crust. Although discontinuous non-uniform stretching models have been successful in explaining some first order features in long-term subsidence patterns of basins, they rely on a number of requirements:

·

The focal depths of earthquakes in old cratons suggest that the upper part of the lithosphere has relatively high strength and shows active seismic activity. In contrast, the underlying lower part is almost completely a-seismic. This difference may contribute to the ductile deformation mechanisms within the lower portion of the lithosphere which may lead to a rheological de-coupling at mid-crustal level, inducing

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M. Cacace · U. Bayer · A.M. Marotta · C. Lempp Figure 3.1.6. Cartoon showing the classical stretching model (“Conventional stretching”) and the alternative (“delayed infill”) model by Van Wees et al. (2000) for the Permo-Triassic evolution of the Southern Permian Basin. Sres depicts the evolution of the relative magnitude of accomodation space as a function of time for both models (red dotted line refers to the conventional stretching model, and black line to the alternative delayed infill model)

·

different amounts of extension within the two layers (e.g., Sibson 1983; Ranalli 1995). However, the existence of such an intra-lithospheric discontinuity is not universally proven. The mechanism by which the lithospheric mantle may stretch differently is an ad hoc requirement which creates another space problem within the mantle.

All the above requirements may be removed by considering a non-uniform but continuous stretching with depth. In such an extension, the stretching is regarded as a continuous function of depth in the mantle-lithosphere, whose rate decreases with depth as the extension is diffused over a wider area. The amount of stretching depends on the depth beneath the crust and on the angle θ between the vertical and the boundary of the stretched region. Greater values of θ increase the amount of the initial subsidence, while at the same time reducing the amount of the post-

rift thermal subsidence. There are two main important implications of stretching the mantle over a wider area than the crust. The first one is that a point located at the rift shoulder experiences an initial uplift followed by a subsequent phase of subsidence. Sleep (1971) has demonstrated that in the absence of erosion it will approximately return to its initial elevation, while if erosion occurs it will sink to a deeper level than the initial one.The second implication is that the predicted stratigraphic onlap at basin margins causes the so-called “steer’s head” geometry during the following post-rift subsidence phase (White and McKenzie 1988).

3.1.2.5 Simple shear model of Wernicke In general, pure-shear models cannot account for the asymmetry and/or uplift of the flanks as often observed in many basins. To overcome these limitations, a differ-

Driving mechanisms for basin formation and evolution

ent method with respect the pure shear approach for lithospheric extension has been proposed by Wernicke (1981, 1985). Based on studies of Basin and Range tectonics, he suggested that lithospheric extension may be accomplished by displacement on a large-scale, gently dipping shear zone which occurs throughout the lithosphere (Fig. 3.1.7). This shear zone transfers extension from the upper crust in one region to the lower crust and lithospheric mantle elsewhere. Consequently, the zone of fault-controlled extension (“thin skinned extensional zone”) and the zone of upwelling asthenosphere (“discrepant zone”) are physically separated. Each of these zones has a different subsidence history. As discussed by Wernicke (1985),

a thin-skinned extensional zone should show tectonic subsidence but no thermal subsidence beneath the extended upper crust. Tectonic subsidence beneath the thin-skinned area is counteracted by tectonic uplift in the region overlying the thinned lower crust and lithospheric mantle, the “discrepant zone”. The “discrepant” zone will then suffer thermal subsidence caused by cooling of the asthenosphere. This discrepant zone should subside to its initial crustal level, or below in the case of erosion, and hence should be the site of a shallow basin which does not present extensional faults, i.e., sag-type basin. Beach et al. (1987) proposed a similar extensional model for the Viking Graben in the northern North Sea. The results ob-

Figure 3.1.7. Normal simple shear model of the entire lithosphere on a low-angle ductile shear zone, after Wernicke (1985). Zone A: region of no extension in both the crust and the lithospheric mantle, i.e. βcrust= βmantle=1. Zone B: region of thin skinned fault-controlled tectonics, i.e. βcrust> βmantle=1. Zone C: region where the shear zone enters the mantle-lithosphere, i.e. (βcrust or βmantle)>1. Zone D: discrepant zone after Wernicke (1985), i.e. βmantle> βcrust=1. Zone E: boundary for the discrepant zone, i.e. (βcrust or βmantle)=1

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tained explained the geometry of the crustal-scale faults but were not able to account for the presence of the thermal subsidence beneath the North Sea. Specifically, the Wernicke shear zone model cannot explain basins which experience a thermal subsidence spatially superimposed on a fault-controlled subsidence.

3.1.2.6 Asymmetrical stretching of the crust In order to overcome the limitations associated with the simple model of Wernicke, Coward (1986) proposed an alternative taking into account the evidence for crustal scale simple shear as well as the subsidence history. The model involves a combination of widespread upper crustal fault controlled extension above a more localised region of concentrated sub-crustal extension (Fig. 3.1.8). Localised extension of the lower crustal/mantle-lithosphere is enforced by strain softening brought about by several causes such as increased geothermal gradient, fabric development, and reduction in grain size. These processes induce a non-uniform deformation in the sub-crustal portion of the lithosphere leading to localised deformation in zones of more intense stretching or simple shear. However, in the upper crust, planar brittle faults tend to lock after a certain amount of extension so that deformation is transferred to new steep faults. According to the model of Coward (1986), the new generation of faults may spread

outwards essentially along one direction away from the initial rift zone thus widening the zone of extended upper crust. Additionally, the presence of some earlier gently dipping anisotropy or compositional layering within the crust may help the expansion of the upper crustal stretching zone. Within this model patterns of uplift and subsidence vary across the basin. A typical sedimentary basin could have an outer zone where only the upper crust is stretched and an inner zone where the upper crust is stretched by a factor, β, smaller than the lower crust/upper mantle, β+β’. The additional stretching factor β’ in the lower portion of the lithosphere is needed to balance the extension in the upper crust at the basin margin. The extension and fault development in the upper portion of the lithosphere is not necessarily symmetrical across the basin. The model prediction may be summarised as follows:

· ·

The zone where only the upper crust has thinned, the “outer zone” of Coward (1986), undergoes an initial tectonic subsidence related to thinning of the upper crust. The amount of this tectonic subsidence is less than in the case when thinning has occurred throughout the entire crustal body. No thermal subsidence will then occur within this portion of the lithosphere. In the “inner zone”, where the whole lithosphere has thinned, both initial tectonic subsidence and later thermal subsidence occur. Stretching of the whole lithosphere by a factor β will cause subsidence: however the additional factor β’ in the lower crust and lithoFigure 3.1.8. Model for het-erogeneous thinning of the lithosphere, after Coward (1986)

Driving mechanisms for basin formation and evolution

·

spheric mantle causes early uplift. Additionally, with extreme density variations in the lower crust a net uplift can occur which raises the stretched upper crust above sea level, thus causing unconformities.

These processes may explain the Late Jurassic-Early Cretaceous “Cimmerian” unconformity observed in the North Sea. According to Coward (1986), the development of Jurassic unconformities in the North Sea may be due to slight uplift caused by heterogeneous stretching during the regional extension rather than being the product of different phases of regional Cimmerian compression.

3.1.2.7 The role of intra-plate stresses: uplift and basin formation in compression Although intra-plate stresses can play an important role in the post-rift evolution of the stratigraphic sequence, early studies of the flexural response of the lithosphere often neglected the contribution of these stresses. In principle, the first-order effects of variations in intra-plate stresses on basin formation and evolution can be estimated within all the above described models for continental deformation. As an example, most of the

Figure 3.1.9. Interpreted structure of the Donbas area along the deep reflection line DOBREflection (top) together with balanced reconstructions of the evolution since the Devonian/Carboniferous boundary (Maystrenko et al. 2003). The dashed line indicates the top of the high velocity lower crust modified after DOBREFRACTION’99 Working Group (2003)

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Figure 3.1.10. Deep Seismic Reflection Profile BASIN9601, illustrated by a line drawing of main reflections superimposed onto the depth converted seismic section, modified after Marotta et al. (2000). Main crustal entities are illustrated. Abbreviations: NHF = Northern Harz Boundary Fault, THZ = Thrust Zone, MH = Moho, HDLC = High Density Lower Crust, NEGB = North East German Basin, Z = base Zechstein reflector

stretching models discussed above can also be applied in compressional regimes causing crustal shortening, uplift or inversion. In the case of McKenzie’s pure shear model, the only requirements are that β<1.0 and that lithosphere is transformed into asthenosphere during subsequent heating of the root. The same of course is true for the other simple non-uniform stretching models. The simple shear model of Wernicke has a compressional counterpart in the inverted southern part of the DnieprDonets palaeorift (Ukraine), the Donbas (Maystrenko et al. 2003) as illustrated in figure. 3.1.9. The basin originated in the Devonian and was induced by a typical wrenching event, well documented by a sequence of mostly small scale basement-faults at the base of the sedimentary sequence. A thermal subsidence phase followed, as expected, with some additional complications in the Permian, not detailed here. Probably at the end of the Cretaceous the Donbas-part of the basin was finally inverted. According to the interpretation of the deep reflection seismic line DOBREflection by Maystrenko et al. (2003), the inversion caused failure at the Moho-level of the old Archaean and cold crust, continuing into a lystric which finally ends at the former eastern end of the basin or rift. In addition a backthrust developed, generating a mega horst or “pop up structure”. Here Wernicke’s model is turned head over heels: it is not the initial stretching event causing the lystric detachment but rather the inversion with a detachment surface below the Moho, an aspect which requires rheological considerations as described in the next sub-chapter. Studies of the tectonic stress field within the lithospheric plates have shown a causal relationship between proc-

esses affecting the plate boundaries and intra-plate deformation (Cloetingh and Kooi 1992). Intra-plate stresses can modulate the long-term thermal subsidence of postrift sedimentary basins (e.g., Cloetingh et al. 1985) and can induce vertical motions of varying directions and magnitudes which are variable along the basin profile. Cloetingh et al. (1992) have demonstrated that for conventional sediment loading models a superimposed compressional stress state causes relative uplift of the basin flanks, subsidence at the centre of the basin and seaward migration of the shoreline, thus giving rise to an unconformity and to the development of an offlapping sequence basinward. In contrast, increasing tensional stresses induce widening of the basin, subsidence of the basin flanks and landward migration of the shoreline together with the development of an onlapping sedimentary sequence. Stress-induced vertical motions of the crust can at the same time influence sedimentation rates. Applying these concepts, a mechanism of compressive-induced flexure of the continental lithosphere has been suggested to explain geological structures for tectonic regional settings where classical models for basin formation have previously failed. The North East German Basin is an example of such a structural setting. The present-day crustal structure below the North East German Basin (NEGB) is characterised by a normal Variscan crust of almost constant 32 km thickness along its margins, thinning to about 22 km under the basin centre (see Fig. 3.1.10) for the seismic cross section. The observed change in thickness of the crystalline crust can be attributed to the change in sedimentary thickness across the basin. However, the observed thinned crust beneath the basin centre lies over an almost perfectly flat Moho which presents a distinct flexural-bulge shape along the southern present-day basin margin. The bulge is located beneath a fault system, (THZ of Fig. 3.1.10), which is interpreted as a thrust zone within the crust and may be related to the Late Cretaceous-Early Tertiary inversion phase. In addition, modelling of wide angle seismic data (e.g., Bayer et al. 1999) and gravimetric studies (e.g., Scheck et al. 1999) have constrained high seismic P-wave velocities and densities in the lowermost crust between Grimmen and the Elbe line (see Fig. 3.1.10 for their locations). A similar feature has also been proposed further to the west on the base of refraction lines crossing the North German Basin (e.g., Thybo 1990; Rabbel et al. 1995). To model the observed structure beneath the NE German Basin, Marotta et al. (2000) developed a simple 2-D thin elastic plate flexural model. In their study, the observed Moho updoming in the southern part of the basin was regarded as a consequence of flexural buckling of the previously thinned lithosphere, induced by the Alpine compressive stress regime. To test this hypothesis, the

Driving mechanisms for basin formation and evolution

Box 3.1.2 Two-dimensional flexure of a thin elastic beam General equation for the deflection of a thin elastic beam

←

flexural resistance

→

←

→ ←

Shear component

→

← →

Isostatic restoring Applied vertical force load

General solution: “Universal flexural profile”

ωb = height of the bulge; wavelength of the bulge; χ0 = point of no flexure;

χ = distance from the origin of the plate;

lithosphere was approximated as a loaded elastic beam by adopting the simple flexural model as often used for the oceanic lithosphere. The effects produced by the compressive forces were modelled as equal to a bending moment applied to one end of the beam plus a vertical load at the free end of the plate in terms of the so called universal flexural profile (see box 3.1.2 for the mathematical formulation). The modelling results supported the hypothesis that the present day Moho topography may be explained as a consequence of flexural buckling of a previously thinned lithosphere under an active compressive stress field which caused subsidence of the North East German Basin during the Late Cretaceous-Early Tertiary inversion event. In fact, for a bulge of wavelength of about 33 km and height of 3 km the predicted and the observed Moho were found to be in very close agreement (Fig. 3.1.11). Moreover, the match between the predicted flexural profile and the

present topography was also satisfactorily explained by considering the eroded amount of sediments originally located above the actual southern margin of the basin. Although the flexural model has some limitations concerning the continental crust due to its complex rheology as discussed below, it does explain well the structure of subducting oceanic plates and in parts the evolution of forearc and molasse basins (e.g., Marotta et al. 2000). Several other studies, focusing on a large number of basins in the northern Atlantic and in the Mediterranean region favour compressional stresses as driving mechanisms for the observed rapid vertical motions (e.g., Cloetingh and Kooi 1992; Reemst et al. 1994). As an example, Van Wees and Cloetingh (1996) applied a three-dimensional flexural model incorporating lateral variations in flexural rigidity and necking depth to quantify the effects of intra-plate stresses on Quaternary accelerated subsidence and uplift in the North Sea Basin and adjacent areas. To explain the short-wavelength Quaternary depocentre in the North Sea,

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Figure 3.1.11. Predictions (red dashed lines) from the flexural study made by Marotta et al. (2000) compared to the interpreted Moho (M) and the Top Cretaceous reflector (TC). TP = Top Palaeozoic reflector, Vo = vertical line load, Mo = bending moment. The post-Zechstein sedimentary fill is also shown by yellow colours

marked by sediment thickness up to 1,000 m, they consider tectonic processes as the main driving mechanism. Their results show that a relative increase of compressive intra-plate forces allows prediction of the accelerated subsidence values up to 700 m, thus successfully explaining the observed Quaternary isopach values. Similarly, Horvath and Cloetingh (1996) were able to explain the overall Quaternary uplift and subsidence in the Pannonian basin by proposing an increase in magnitude of horizontal compressional intra-plate stress in their two dimensional flexural model. A more sophisticated model was proposed by Grünthal and Stromeyer (1992, 1994). In their study, they applied a 2-D, steady-state, elastic finite-element approach to model the observed direction of the maximum compressive horizontal stress, SHmax, for the brittle domain of the crust for the entire western portion of the Eurasian plate. In order to model the present-day tectonic setting as well as the related style of deformation observed in Europe, Grünthal and Stromeyer (1992, 1994) tried to couple lateral local variations of lithospheric stiffness with specific dynamical boundary conditions. The plate tectonic scheme adopted to simulate the main forces acting at the boundaries of the Eurasian platform comprised: (i) Ridge-push forces from the central and northern segments of the Mid-Atlantic; (ii) continental collisional forces arising from the northward-directed motions of Africa relative to Europe; and (iii) push forces acting across the Anatolian and Aegean microplates. Lateral stiffness contrasts were introduced in terms of variations in rigidity values, i.e., Young’s elastic modulus parameter, applied to different lithospheric blocks, as well as in terms of local extensional and shear zones. The lateral heterogeneities in the elastic properties considered by Grünthal and Stromeyer (1994) were introduced in order to reproduce small scale features such

as the Adriatic Promontory, the Bohemian Massif, and the Pannonian Basin. To constrain the modelling results a step-wise trial-and-error approach was adopted taking as test criterion the direction of the maximum horizontal stress component (SHmax) based on stress data interpolation. The fit between the generalised trajectories and the derived SHmax direction was found remarkably. The calculated horizontal crustal stress orientation reproduced the observed broad-scale pattern, reflecting an almost uniform and consistent NW-SE direction of maximum compression in the western domain of the study area and a gradual bending toward NE-SW approaching the eastern part of the Eurasian plate. Moreover, the introduction of small-scale features, i.e., sub-domains with different elastic stiffness, made it possible to resemble more regional features like the fan-like crustal stress pattern along the western Alpine sector and the observed NE-SW SHmax direction in the Dinarides area along the south-eastern margin of the European plate. The main simplification common to all the previously described models was to adopt an elastic plate model. Plate interiors may retain an elastic “core” under very low deviatoric stresses acting on relative short time scales. As the magnitude of such deviatoric stresses increases, frictional or brittle failure mechanisms will be set up within the crust. Consequently, the elastic approximation prevents a quantification of the role played by the different driving mechanisms which are active during continental deformation. In this context, elasticity tends to overestimate the influence exerted by boundary and/or internal loads on the stress field and deformation patterns and is not able to correctly account for more local features related to the internal thermal and mechanical, i.e., rheological and compositional, structure of the lithosphere. At the same time, the incorporation of weak and/or stiff areas in terms of lateral variations of not-well constrained and unrealistic elastic parameters cannot help in understanding the nature, whether it is thermal, mechanical/kinematic or compositional, of the driving mechanism of the observed local anomalies. In addition the assumption of an elastic rheology is incompatible with the occurrence of intra-plate seismicity (e.g., Cloetingh et al. 2006) and may also lead to large modelling errors (e.g., Bird 1999).

3.1.3 Rheological models 3.1.3.1 The role of rheology on the modes of continental deformation Continental deformation is mainly determined by the rheological structure of the lithospheric plate. The rheology

Driving mechanisms for basin formation and evolution

of the lithosphere is a function of several controlling factors: (i) its composition and structure, (ii) pressure, (iii) temperature, (iv) principal deformation mechanisms at given P-T conditions, and (v) state of stress. Variations of these parameters with depth may lead to a so-called “sandwich-like”, i.e., stratified, structure with an alternation of rheologically weak layers and relative stronger layers. As derived from experimentally constrained results, it is generally assumed that the deformation regime for any type of rock may be broadly subdivided into two domains: a brittle (frictional) domain within the colder and/or upper portions of the lithosphere, and a ductile domain at deeper levels and/or higher temperatures (higher than about one-half of the solidus temperature of the relevant material).

The distinction between brittle and ductile domains is not easy to achieve. However, as a first order approximation, the distribution of crustal and upper mantle seismicity outside subduction zones can be taken as an indication of the depth extent of brittle layers. Although some localised diffuse deformation, i.e., pressure solution transfer, may occur at relatively low temperature conditions, in the brittle domain rock failure is predominantly governed by the Coulomb-Navier shear failure criterion, usually referred to as Byerlee’s law (see Fig. 3.1.12). According to Byerlee’s law, friction on rock surface is a non-linear function of the overburden pressure, strain and sliding velocity (Behn et al. 2002) while it is relatively unaffected by temperature variations. On the other hand, ductile deformation is governed by a temperature-dependent power law creep equation between stress and strain rates. In gen-

Figure 3.1.12. Examples of strength envelopes, modified after Marotta et al. (2000). A Stress distribution for a multi-layered continental lithosphere. Brittle failure, red solid curves, as described by Byerlee’s law: σB= (σ1 - σ3) B = β · z. It increases linearly with depth and pressure. Here, σB refers to the brittle yield strength, σ1 and σ3 are respectively the maximum and the minimum principal stress, Ζ is the depth, and β represent the brittle failure coefficient. It is dependent on the type of fault, the angle of fracture and the pore pressure and its values amount to 16 MPa km-1 for extensional settings and 40 MPa km-1 for compressional settings (e.g. Lynch and Morgan 1987; Jimenéz-Munt et al. 2005). Ductile behaviour, blue dashed curves, is simulating adopting a steady state power law creep relationship between stress and strain rate: σD= (σ1 - σ3) D =(ἑ/ἑo)Ine(E/nRT). ἑ is the effective strain rate given by the second invariant of the strain rate tensor, R is the gas constant, T is the absolute temperature, and ἑo, n , E are material constants depending on the rock type. The steady state power law flow creep equation indicates an exponentially decreasing of critical differential stresses with increasing temperature. The mantle and the crust are assumed decoupled. Grey dash-dotted lines indicate the elastic bending stress, σE, for both the crust and the mantle-lithosphere. It is a function of the local curvature. The bottom of the mechanically strong crust and mantle are indicated by h1 and h2 respectively. For each layer, zni (i=1, 2), represents the depth of non-flexural strain (i.e. depth of necking after Braun and Beaumont 1989a,b). B Final yield strength envelopes, σy with σy =± min{| σB | ; | σD |}, as represented by the solid red contour function. It is positive for extension and negative for compression. Comparing the yield strength envelopes, σy, with the elastic stress (dash-dotted lines), σE it is possible to obtain the differential stress for a given strain rate and curvature. Different colours are used to enlighten the elastic (light green) and the brittle or ductile domain (yellow). When the elastic bending stress crosses the yielding curve the rocks fail and become brittle or ductile

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eral, before reaching the condition of steady-state creep, i.e., strain rate constant under constant stress and temperature, ductile flow is governed by instantaneous elastic deformation and transient creep, i.e., strain rate decreasing with increasing time. However, with the exception of few geodynamical processes, e.g., post glacial rebound, the elastic and transient strains may be regarded as negligible with regard to steady-state strains in tectonic deformation. The first order rheological behaviour, brittle or ductile, at any given depth is determined by the relative (minor) magnitude of frictional and creep strength. If the critical strength for frictional sliding is less than that required for

ductile behaviour, failure dominated by frictional sliding will occur. In the opposite situation, ductile flow will be the dominant mechanism of deformation. The so-called “brittle-to-ductile transitions” correspond to the depth levels associated with the equality of frictional and creep strength for a given strain rate value. Vertical integration of the strength of the lithosphere at different depths allows the evaluation of the total strength of the lithosphere (England 1983). Within basin modelling the rheology is usually approximated by a finite number of lithological units of uniform structural composition. The depth-dependency of the rheological structure of continental litho-

Figure 3.1.13. Rheological profiles for different types of lithosphere composition and structure, modified after Cloetingh and Burov (1996). Depending on the adopted rheology, mechanical decoupling zones may appear at different levels. A.Simple case of a uniform, olivine dominated, lithosphere. B.Continental lithosphere consisting of a crust, quartz dominated, and of an underlying mantle, olivine dominated. C.Continental lithosphere consisting of an upper crust (quartz), a lower crust (diorite) and a lithosphere mantle (olivine). Different colours in each panels illustrates the effects of variations of strain rates on the strength of the lithosphere

Driving mechanisms for basin formation and evolution

sphere is addressed by the construction of strength envelopes, i.e., rheological profiles, as extrapolated from rock mechanics data and additional constraints such as seismicity, gravity and petrology. The concept of strength envelopes was firstly developed by Goetze and Evans (1979) and it is nowadays well known. Strength envelopes allow the identification of brittle and ductile layers and the estimation of lithospheric strength. Essentially, they consist of two different types of curves. The straight lines represent brittle failure. Following Byerlee’s law they show increasing rock strength with increasing depth. The curve lines stand for ductile deformation. The strength described by such lines decreases exponentially with depth. Each line is for a given strain rate that is assumed constant throughout the entire lithosphere. Combining the rheological laws describing the different modes of continental deformation it is possible to describe a piece-wise continuous contour function given by the less between the two contributions above mentioned, i.e., brittle or ductile. Such a curve defines the strength envelope, (see Fig. 3.1.12). Generally speaking, the lithosphere may be regarded as consisting of two strong layers: the middle brittle crust and the uppermost brittle/ductile mantle. However, the rheological layering of the lithosphere is much more complex than usually assumed while constructing strength envelopes so that several assumptions regarding deformation style are very common. Figure 3.1.13 shows strength profiles related to different thermal regimes as well as different structures and compositions of the lithosphere. The firstorder predictions of rheological studies may be confirmed by geophysical observation. However, the distribution of deformation within the lithosphere is highly inhomogeneous and this may induce large uncertainties and errors in constructing rheological profiles. In this regard, many studies (e.g., Roth and Fleckenstein 2001; Marotta et al. 2002; Maystrenko et al. 2006) have demonstrated that the growth of salt structures as represented by salt pillows or diapirs within the sedimentary sequence strongly influences basin evolution. In fact, local salt movements may affect and even determine the stress pattern in the overburden (e.g., Bayer et al. 1999; Marotta et al. 2001; Röckel and Lempp 2003; Lempp and Lerche 2006). This decoupling role played by salt structures is demonstrated by the North German Basin (NGB), along the southern margin of the Central European Basin. With respect to the basin dynamics, the NGB underwent a NS-oriented compression and an EW-oriented extension as a result of the plate tectonic situation between the Mid-Atlantic Ridge across its western side and the Alpine orogenic belt in the south. However, especially in the suprasalt layers (above the Zechstein salt deposits), EW compression and NS extension are dominant in combination with lithological-dependent stress gradients leading to a highly in-homogeneous suprasalt stress field. The

underlying evaporitic sequences of Zechstein age, with their widespread and great thickness, decouple the subsalt stress field from the suprasalt. The uprising salt masses, with generation of salt pillows and/or diapirs, influence the stress transfer conditions between suprasalt and subsalt domains especially in regions with reduced salt thickness. Moreover, tectonic faults connecting different depths additionally influence the state of stress in the suprasalt sequence. This results in a state of stress in the sediments which differs significantly between the subsalt and the suprasalt layers. In the subsalt layer a more uniform orientation of the maximum horizontal stress component, an approximate NS direction being dominant. Moreover, a nearly constant gradient of the minimum horizontal stress is also present. Contrastingly, within the suprasalt sequences the orientation as well as the gradients of the detectable stress components are more scattered and vary more on regional scale. Consequently, the stress field in the suprasalt is inhomogeneous in comparison with that in the subsalt stress field as shown in figure 3.1.14. In addition, within the suprasalt the gradients of detectable stress components are more scattered and vary more on a regional scale (Röckel and Lempp 2003).

3.1.3.2 Limitations of a kinematic approach to continental deformation The fundamental attribute of all the above described kinematic models for basin formation and evolution is that the deformation is imposed by prescribing a velocity field without incorporating any constitutive equations. However, due to the high variability of deformation modes that can be prescribed by prior imposed velocity fields, such models have been able to describe the main features of basin evolution. The major limitation related to a kinematic approach is the inability to establish a causal relationship between the imposed deformation pattern and the mechanical behaviour of the lithosphere. Many recent studies have tried to overcome this limitation by introducing realistic rheological constraints to their model formulations in order to control and explain the imposed mode of deformation. An example is given by Marotta et al. (2000). To constrain the results obtained from the flexural model for the North German Basin, discussed above, they performed an independent rheological study along the seismic reflection profile shown in figure 3.1.10. To model the rheological structure of the lithosphere under the basin, an elasto-plastic rheology at the time of loading was adopted within two different rheological models. In the first case they considered the crust composed of two structural layers, sediments and one single crystalline crustal body, while in the second model they further subdivided the crust into two different sub-layers of varying thicknesses in accordance with the seismic reflection

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M. Cacace · U. Bayer · A.M. Marotta · C. Lempp Figure 3.1.14. (a) SHmax direction in subsalt layer of the NGB after Röckel and Lempp (2003). Different colours refer to different authors. (b) SHmax direction in suprasalt layer of the NGB after Röckel and Lempp (2003). Different colours refer to different author. (c) SHmin and Svert gradients in suprasalt and subsalt domains, after Röckel and Lempp (2003)

profile of figure 3.1.10. Brittle failure was modelled assuming plastic rock behaviour, i.e., Byerlee’s law, and ductile behaviour was implemented through the use of a steady-state power law creep equation. For both the rheological modelling analogues, the calculations of rheological profiles and total elastic thickness of the lithosphere were in agreement with the results obtained by the simple elastic model. In fact, the elastic thickness predicted by the rheological study resembled the one derived from the flexural elastic model confirming the presence of a weak lithosphere along the southern margin of the basin and a probable decoupling between the lower and the upper lithosphere at the Moho level as already suggested by the outcomes of the elastic model. The use of rheological constraints in kinematic models has lead to some improvements in understanding basin formation and evolution. However, in order to account for continental

deformation dynamic modelling approaches are necessary. Only dynamic models, explicitly invoking constitutive relations which relate dynamic quantities, i.e., stresses, with kinematic quantities, strain and/or strain rates, by the mean of realistic material parameters, can provide insight into the physical processes underlying basin evolution.

3.1.3.3 Dynamic models for basin formation and evolution Although intra-plate stresses have been shown to be very important in controlling the subsidence record as well as the stratigraphic architecture of extensional basins, most models for basin formation and evolution, e.g., simple and/or pure shear models, have usually related lithospheric strain patterns to unspecified and/or unrealistic stress fields. In a similar way, models for basins developed under com-

Driving mechanisms for basin formation and evolution

pressional tectonic settings have been mainly based on lithospheric flexure profiles without addressing the role of the active compressional tectonic stress field. Additionally, results from different studies, e.g., Cloetingh et al. (2006), have demonstrated that lateral strength variations of intra-plate lithosphere are primarily caused by variations in the mechanical strength of the lithospheric mantle which are related to variations in the thermal structure of the lithosphere (e.g., Goes et al. 2000b). To overcome these restrictions, attention was focused on dynamic modelling techniques. First attempts to simulate anelastic deformation within the lithosphere restricted the deformation to plane horizontal strains in a vertical cross section perpendicular to the strikes of linear structures, i.e., the plane strain approach. A measure of qualitative success has been obtained in modelling convergent tectonic scenarios (e.g., Tapponier and Molnar 1977), but the assumptions of this approach limit quantitative applications to continental deformation. The major limitation of this method is the restriction to processes that can be approximated as purely two dimensional. An alternative approach to

model continental deformation is the “plane stress” model, now widely referred to as the thin-sheet or thin-plate model (e.g., Bird and Piper 1980; England and McKenzie 1982, 1983), see box 3.1.3 for its mathematical formulation. In this technique, deviatoric stresses are assumed to vanish beneath the lithosphere, whereas within the lithosphere the differential equations for the equilibrium of stresses are vertically integrated following the isostatic approximation, i.e., lithostatic vertical stress condition. The equations of conservation of momentum and mass are numerically solved, within the finite difference and/or finite element techniques, in order to predict (horizontal) deformation velocities, stresses, and therefore strain and strain rate patterns. Once the dynamic quantity, i.e., stress, is obtained it is possible to derive the kinematic quantity, i.e., strain. The quality of numerical models can then be tested and quantified by comparing the obtained results with observations. The thin-sheet approach has been widely used by several research groups which have applied it to a variety of

Box 3.1.3 Mathematical formulation of the thin-sheet model

Navier-Stokes equation for geological time scales:

Model assumptions

Spherical coordinates

η = vertically averaged viscosity of the lithosphere; hc = crustal thickness; L = lithospheric thickness; ρc = density of the crust; ρm = lithospheric mantle density; R = Earth’s radius

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tectonic scenarios. Although an accurate and more realistic representation of lithospheric strength should involve a fully three-dimensional integral model, the thin-sheet approach can be regarded as a good and reasonable approximation to continental deformation. At the same time, it entails lower computational costs and elaboration times than three-dimensional integral models. Previous thinsheet models accounted only for mechanical deformation (e.g., England and McKenzie 1983). Further implementations included temperature-dependent creep parameters as well as the possibility to address more complex processes like kinematic detachments within the crust and the mantle-lithosphere (Bird 1989; Sonder and England 1989). In the meantime, several numerical studies also started to address lithospheric deformation as a function of internal structure and composition. Within this context, many groups of modellers have tried to represent lithospheric strength by use of a single vertical averaged power law rheology (e.g., Vilotte et al. 1982, 1984; England and McKenzie 1983; England and Houseman 1985; Houseman and England 1986a,b). They have found that for a typical two-layered lithosphere, crust and mantle, the vertical averaged stress-strain rate relationship can be approximated over a wide range of strain rate values (10-13-10-17 s-1) by a power law rheology which accounts for more than one deformation mechanism. Using this approach, only two parameters are necessary to describe the mechanical behaviour of continental lithosphere: the stress exponent and the stress coefficient. The applicability of this simple approach is limited by the strong dependence of these two parameters on the geothermal field in the uppermost mantle as well as on the state of stress within the upper crust (e.g., Houseman and England 1986a,b; Sonder and England 1989). Other authors in the past have used a more complex combination of plastic and viscous rheologies (e.g., Villotte et al. 1982; Bird 1999). Such an approach reflects the concept that friction on faults and plastic failure in the upper portions of the lithosphere may control the vertical averaged rheology of the entire continental lithosphere. The main stumbling block is to determine a single vertical averaged rheology on the base of the distribution of deformation mechanisms with depth, this requiring a very accurate knowledge of the thermo-mechanical structure of the lithosphere. Rheological stratification of continental lithosphere and variation in the local geothermal field may account for the observed heterogeneous style of continental deformation (e.g., Behn et al. 2002). Following these principles many authors in the past have tried to address the occurrence of weak and/or stiff domains in terms of lateral variations of the thermal regime (e.g., Molnar and Tapponier 1981; England and Houseman 1989; Tommasi et al. 1995). However, the lack of a full thermomechanical formulation has strongly limited the ability

to understand the retro-effects between deformation and temperature variations. More recent versions have been implemented where lateral variations of lithospheric stiffness have been coupled with specific boundary conditions in order to model the observed style of deformation in both collisional and extensional tectonic setting (e.g., England and Houseman 1985; Vilotte et al. 1986; Lynch and Morgan 1990; Tommasi et al. 1995; Marotta et al. 2001; Jimenéz-Munt et al. 2001). In this context, Marotta et al. (2004) applied a spherical viscous thin-sheet model to investigate the pattern of crustal deformation within Central Europe as a function of tectonic loading forces and changes in the internal structure of the lithosphere. They modelled the main tectonic forces acting across the boundaries of the European plate (i.e., Mid-Atlantic Ridge push forces, collisional forces related to the collision between Africa and Europe). At the same time, they modelled lateral variations in lithospheric stiffness imposing different viscosity values to the subdomains into which the entire study area was subdivided. To quantify the sensitivity of the modelling results with regard to changes of the imposed boundary loads as well as to lateral variations in lithospheric stiffness, the predictions derived from the tectonic model were compared first to the deformation pattern computed for the same region using a spherically symmetric, self-gravitating, visco-elastic (Maxwell) Earth model of Glacial Isostatic Adjustment (GIA) and finally to observed baseline rates (i.e., length changes) relative to three different fixed sites within the permanent ITRF2000 and BIFROST GPS networks. The reference sites (Potsdam in Germany, Vaas in Finland and Onsala in Sweden) were chosen as locations with different latitudes in order to resemble the different levels of deformation associated with tectonics and GIA. The model demonstrated that tectonic deformation associated with boundary stress forces (i.e., Middle Atlantic Ridge push forces and African to Europe collision-related forces), GIA effects, and lateral variations in the strength of the lithosphere must be taken into account simultaneous­ ly to reconcile the broad style of intra-plate deformation in Europe. In fact, with reference to the ITRF2000 velocity solution database, the best fit with observations was achieved combining both, the tectonic predictions comprising a weakened lithosphere in the Mediterranean subdomain together with the standard GIA predictions. The study described above is a regional analysis concerning the broad scale stress and/or strain field. It covers a very wide region extending from the Mid-Atlantic Ridge to about 40 °E, and from the natural plate boundary between the African and the Eurasian plate to about 90 °N. The great extension of the study area allows prediction of possible changes in the stress field and in lithospheric deformation style within a minimum wavelength

Driving mechanisms for basin formation and evolution

of a few hundreds of kilometres. Consequently, when the modelling results are compared with available datasets (e.g., the World Stress Map database and/or GPS detected crustal velocities) a satisfactory agreement is achieved only for first order features. Many other large-scale models have been carried out in the past (e.g., Grünthal and Stromeyer 1994; Richardson and Coblentz 1994; Gölke and Coblentz 1996). The main problem was the lack of detailed information on the structure of the lithosphere. This lack of knowledge led to the adoption of unrealistic simple rheologies thus impeding a full examination on the dynamic aspects of continental deformation, and at the same time preventing basin-wide models. However, substantial progress has been made during the last decade in terms of availability of new and high-quality data sets both from well investigated areas and from new unexplored areas. This new information has shed new light on the subsidence and uplift history, the temporal and spatial evolution of the stress and strain fields, the thermal history and structural development for several basin systems. One of the best examples is the Central European Basin System (CEBS) in northern Europe, (see Fig. 3.1.15) for its location and main features. Within the CEBS several deep reflection and refraction seismic experiments (e.g., BIRPS, DEKORP, EGT, DOBRE, BABEL, MON-LISA), gravity modelling (e.g., Yegorowa et al. 2007), seismic to-

mographic studies (e.g., Gregersen et al. 2002) as well as structural crustal modelling (Gemmer and Nielsen 2001; Scheck-Wenderoth and Lamarche 2005) have been carried out during the last decade. The obtained results have elucidated the structural configuration of the lithosphere beneath the basin system and have provided a deeper understanding of the geodynamic processes which govern lithospheric deformation. Based on these advancements, new dynamic models have been developed which successfully incorporate realistic constitutive equations in the lithosphere. This allows the simulation of more complicated rheologies including structural and/or thermal inhomogeneities of earth’s materials as well as testing of their role in basin development under different tectonic loading boundary conditions. A first and relatively simple example is given by Marotta et al. (2002). In their study, they adopted a viscous thinsheet model to investigate the style of crustal deformation in the North German Basin (NGB) in northern Europe. The NGB has a complex dynamic setting. The presentday regional stress field shows a broad-scale NW-SE direction of maximum horizontal compression in the western part of the basin with a clockwise bending toward N-S to NE-SW moving further to its eastern flank (e.g., Roth and Fleckenstein 2001; Reinecker et al. 2005). This fan-

Figure 3.1.15. Location of the Central European Basin System (dotted thick grey line), modified after: Ziegler 1990; Pharaoh 1999; Bayer et al. 2002; Scheck et al. 2003a,b. The main structures which framed the CEBS are: (i) the Tornquist Zone consisting of the Sorgenfrei-Tornquist-Zone (STZ) and the Teisseyrie-Tornquist-Zone (TTZ); (ii) the Elbe Fault System (EFS); (iii) the Ringkøbing-Fyn-High (RFH); and (iv) the Elbe Lineament (EL). Also shown are the three main Permian basins (the North German Basin, the Danish Basin and the Polish Trough), the major grabens developed during the early phases of the Mesozoic: the Central Graben, the Horn Graben, the Glückstadt Graben

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like opening in the stress direction is a feature which occurs also in other basins, e.g., the Pannonian Basin, and the Polish Trough (e.g., Jarosinski 2005). In addition, new sets of borehole stress data, (Roth et al. 1999; Roth and Fleckenstein 2001), show a pattern of SHmax direction and orientation decoupled at the level of the Zechstein salt deposit. As previously described, while the suprasalt data show an almost random pattern ranging between NW-SE to W-E orientations, below the Zechstein salt layer the stress orientation changes into N-NE. Several models have been previously proposed to resemble the presentday stress field under the NGB. The majority of them have tried to relate the observed stress field pattern just to effects connected with plate boundary forces and not well specified intra-plate stress sources (e.g., Richardson 1992; Grünthal and Stromeyer 1994; Richardson and Coblentz 1994; Gölke and Coblentz 1996). The results have mostly failed to explain the observed local deviation in the stress direction under the NGB.

Figure 3.1.16. (a) Finite element mesh for the study area of Marotta et al. (2002). The elements in which the local effective viscosity was changed (two orders of magnitude higher than the surroundings) to resemble a stiffer lithosphere under the North German Basin (NGB) are painted in red colours. (b) Assumed boundary conditions, given in terms of velocity (ui, i=1,2), shear stress (τ12), and crustal thickness (s), used in the study of Marotta et al. (2002). On the northern boundary of the study area (CDF in the upper panel) no slip conditions (ui=0, i=1,2) and constant crustal thickness (s=s0) were applied. At the lower boundary (resembling the Alpine Front) both crustal thickness and horizontal velocity, after Schmid et al. (1996), were held fixed. At the lateral boundaries a constant crustal thickness (s=s0) and zero shear stress (τ12=0) were applied. Two different models have been tried in terms of velocity boundary conditions along the lateral sides: the first with closed lateral boundaries (zero horizontal velocity components), and the second with open lateral boundaries and continuity of the horizontal component of the velocity. Abbreviations: CDF=Caledonian Deformation Front, HG=Horn Graben, NGB=North German Basin

To overcome this problem as well as to resolve the details of the local perturbations of the stress field around the NGB, Marotta et al. (2002) tried to couple both tectonic boundary forces and lateral inhomogeneities in the internal structure of the lithosphere. Following previous studies which supported the idea of a causal relation between the present-day basin structure and the Late Cretaceous-Early Tertiary Alpine compression-induced deformation (e.g., Bayer et al. 1999; Marotta et al. 2000; DEKORP BASIN Research Group 1999), they focused on the effects of the Alpine orogen-related, transmitted compressive forces on the local stress field. Lateral variations in the internal

Figure 3.1.17. Modelled SHmax direction as obtained by Marotta et al. (2002). As a consequence of considering an increased lithosphere under the North German Basin the direction of SHmax undergoes a local rotation toward the external part of the basin

Driving mechanisms for basin formation and evolution

structure of the lithosphere were then inserted increasing the strength of the lithosphere below the NGB (about two orders of magnitude higher than the surrounding areas), (see Fig. 3.1.16). A relatively stronger lithosphere under the NGB was chosen in accordance with results from various studies (e.g., Thybo 1990; Rabbel et al. 1995; Bleibinhaus et al. 1999; Scheck et al. 1999; Bayer et al. 1999; Cloetingh et al. 2006), suggesting a similar consistent trend for other geophysical parameters: (i) observed high lower crustal seismic velocities as derived from refraction data, (ii) high density values in the lower crust as indicated by gravity modelling, and (iii) observed local a-seismicity under the NGB with respect to the surrounding areas. The main results using a strong lithosphere below the NGB were that the basin underwent almost no crustal deformation, and by behaving as a “barrier”, the stiffer basin prevented the release of (compressive) stresses. Consequently, this frame determined a local rotation toward the external parts of the basin in the direction of the regional stress field resembling the observed fan-like stress opening as shown in figure 3.1.17.

they adopted a steady-state, an-elastic thin-plate model (the program “Shell” developed by Bird and Kong 1994) which provides frictional/dislocation creep rheologies, spherical shell (finite) elements and integrations of faults. Intra-plate deformation was investigated by varying rheologies and boundary conditions (GPS measured residual velocities).The main tectonic forces simulated were those related to the opening of the Atlantic (i.e., ridge push forces), the Alpine orogeny (i.e., collisional forces), and the horizontal forces of the postglacial (Late Pleistocene) isostatic rebound of the Scandinavian Shield (e.g., Scherneck et al. 1998). The effects of topography were also investigated. The role played by the presence of weak domains and their reactivation on intra-plate deformation,

Although based on a very simplified structure of the lithosphere (linear creep rheology), the results of the model by Marotta et al. (2002) suggests that local heterogeneities in the rheological configuration of the lithosphere may be a key factor for the distribution of (compressive) stresses. In general, dynamic models require an initial perturbation (i.e., heterogeneity) in order to focus deformation within a finite area. The thin-sheet approach can incorporate initial perturbations in order to modify the mode of deformation. Depending on their nature, lateral inhomogeneities may give rise to different (lateral) strength variations: thermal-, rheological- or fault-related. In principle, deep faults can concentrate deformation and modify the resulting deformation pattern because of the associated mechanical and thermal effects. However, being by definition surfaces of discontinuity, faults and/or detachments cannot be realistically incorporated in continuum models. In addition their position and orientation have to be determined a priori. To overcome these problems many numerical studies have tried to built fault zones into the models by means of ad hoc anisotropic mechanical properties (e.g., plastic shear bands or slippery nodes) or elongated areas of low viscosity (e.g., Braun and Beaumont 1989a,b; Melosh and Williams 1989; Dunbar and Sawyer 1989). An example is the study of Kaiser et al. (2005). In order to understand the recent structural evolution of the North German Basin, they performed a finite elementbased dynamic modelling study of the present-day stress condition and distribution within the basin. In their study

Figure 3.1.18. (a) Finite element grid and modelled faults (thick red coloured lines), modified after Kaiser et al. (2005). (b) Model results (SHmax directions and stress regime) from the study of Kaiser et al. (2005)

was addressed by Kaiser et al. (2005) including within their model formulation a set of (22) deep-seated basement faults together with prominent fault zones derived from different sources as shown in figure 3.1.18a. Local faults and faulted salt migration structures, i.e., salt diapirs, were not taken into account. Their results, (see Fig. 3.1.18b), resemble the observed direction of the principal regional horizontal stress field. At the same time, the results predicted major deviations of stresses occurring along strong contrasts in the lithospheric structure (i.e., faults) which significantly affected the stress pattern. The local bending in the principal horizontal stress direction under the area of the North German Basin was shown to be a function of tectonic loading forces (mainly those related to the postglacial rebound) and major lithospheric structural contrasts. The results of Kaiser et al. (2005) suggest that the interplay of stresses and inherited tectonic features may be related to weak zones (fault zones in their model) which

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are permanently prone to reactivation and can have a strong impact on the deformation pattern. Moreover, observations from old basin systems characterised by multiple phases of extension and inversion suggest that lithospheric deformation is markedly influenced by preexisting heterogeneities in the lithosphere. Consequently, continental deformation does not show an almost uniform pattern but tends to remain localised under rheologically weak areas for time-spans long enough to involve reactivation of tectonic structures at different scales, e.g., from basement-scale reactivations to large lithospheric-scale reactivations. The evolution of complex basin system is strongly influenced by repeated reactivations of such preexisting crustal and deeper discontinuities both during extensional and compressional tectonic scenarios. The widespread presence of faults and shear zones within continental lithosphere has led to the idea that continental deformation might be mainly localised at pre-existing faults. The main point is that faults may be regarded as the weakest zones thus most prone to deformation with first-order influence on intra-plate stresses.

However, this deformation style contrasts with more general cases in which deformation re-works larger volumes of lithosphere without requiring complete reactivation of individual faults and/or shear zones. This feature is illustrated by the process of basin inversion. However, it is very difficult to establish whether or not a fault has been reactivated or to predict why some faults are reactivated and others not. One possible explanation of this may be that upper crustal processes might be driven by deeper lower crust and/or mantle processes and that, consequently, basement reactivation may be regarded as a shallower expression of deeper heterogeneities. Inherited material (i.e., rheological) heterogeneities and thermal effects may strongly control subsequent deformation stages of basin evolution. The Central European Basin (CEBS) is one of the most suitable “natural laboratories” to apply these dynamic concepts. Several studies (e.g., Vejbaek 1997; Berthelsen 1998; Scheck et al. 2002a) have suggested that basin formation and evolution was mainly controlled by the presence of deep-reaching zones of lithospheric weakness. Moreover, other studies (e.g., Bayer et al. 1997) have related these zones of observed stress and strain loFigure 3.1.19. Base maps used in the study of Cacace et al. (2007) to represent the thickness variations of the different lithospheric structural layers. (a) Sedimentary layers modified after Scheck-Wenderoth and Lamarche (2005). The main depocentres shown are (1) NWSE-trending with maxima under the Permian Basins: the North German Basin (NGB), The Norwegian-Danish Basin (NDB), and the Polish Trough (PT), and (2) N-S-oriented below the main Mesozoic grabens: Central Graben (CG), Glückstadt Graben (GG) and the Horn Graben (HG). (b) Crustal thickness variations modified after the European Moho base-map published by Ziegler and Dèzes (2006). It shows sev­ eral variations in the thickness of the crust ranging from 25 km in the central part down to 48-50 km in the easternmost part (EEC domain). (c) Lithosphere-asthenosphere isothermal boundary as derived from the TOR seismic experiments results (Gregersen et al. 2002). Two sharp and steep transitions are illustrated: the first one at the northern rim of the Tornquist Zone (STZ) near the border between Sweden and Denmark with depths of 200250 km, and the second one near the southern edge of the Ringkøbing-Fyn-High (RFH) with smaller differences in the depth of the lithosphere (from 90-100 km of depth to 120150 km). Abbreviations: EEC=East European Craton; STZ=Sorgenfrey Tornquist Zone; TTZ=Teysserie Tornquist Zone; PT=Polish Trough; RFH=Ringkøbing-Fyn-High; NDB=Norwegian-Danish Basin; NGB=North German Basin; CG=Central Graben; GG=Glückstadt Graben; HG=Horn Graben

Driving mechanisms for basin formation and evolution

calisation to the underlying structure of the lower crust or the lithospheric mantle, where areas of reduced viscosity or elevated temperature can play a key role in focusing deformation. In addition, the evolution of the basin system was also strongly influenced by the growth of salt structures (e.g., Maystrenko et al. 2005b). This is reflected by the lack of deep-reaching faults within the CEBS. In contrast, many faults detach at the salt level, not reaching the basement (e.g., Scheck et al. 2002a). Based on these results, Cacace et al. (2007) presented an integral study investigating the influence of deep crustal and shallow mantle structures and their related rheologies on the recent stress and strain patterns within the CEBS. The distribution of stresses and strains within the basin system was investigated through the use of a suit­able spherical, finite element, thin-sheet model. The impact of variations of compositional parameters, thermal perturbations and variations in both crustal and lithospheric mantle thickness on the setting of the regional stress field and deformation style was quantified by a preliminary independent thermo-mechanical study of the area under investigation. The feedback effects between deforma­tion and temperature variations were investigated through the use of a 3-dimensional finite-element model to solve the heat conduction energy equation. The solution for the temperature field with depth took into account all the most important rock properties (i.e., radiogenic heat production within the crust, non-constant thermal conductivities, thermal blanketing of the sediment fill). Depth-related effects (i.e., crustal and mantle thickness variations) as well as effects due to the rheological stratification of the lithosphere were also addressed. The lithospheric plate was

considered as a set of four structural layers representing: (i) the sediments, (ii) the upper and (iii) lower crust, and the (iv) mantle-lithosphere, (see Fig. 3.1.19) for the respective base maps. Brittle-plastic (Byerlee’s law) or ductile (steady state power law creep flow or Dorn’s law) was implemented for each layer depending on the local thermal state and strain rate. Lateral structural contrasts were addressed by grouping the entire area into different lateral sub-domains following the constraints given by previous large scale seismic studies (e.g., Eugeno-S Working Group 1988; BABEL Working Group 1993; MONA LISA Working Group 1997a,b; DEKORP-BASIN Research Group 1999; Guterch et al. 1999; Scheck et al. 1999; Yegorowa et al. 2007) as shown in Fig. 3.1.20. The consistency of the model was constrained by direct comparison of the modelling results and two independent set of data, these being the present-day regional stress field obtained from the “World Stress Map Project” as shown in figure 3.1.21 (Reinecker et al. 2005) and the strain rate eigenvectors derived by the last ten year GPS observations of the ITRF2000 velocity solution database (Altamini et al. 2002), see Fig. 3.1.22. To determine the most suitable boundary conditions, the tectonic loading scheme used in the model of Marotta et al. (2004) was adopted adjusting the final grid shown in Fig. 3.1.20. The results so obtained demonstrated that the observed deformation and stress patterns in the CEBS can be reproduced with reasonable agreement by the model, regarding both the overall trend as well as more local features. At the same time, the results strongly suggested that continental deformation is localised by initial thermal, mechanical and compositional heterogeneities in the lithosphere. AcFigure 3.1.20. Finite element grid used in the study of Cacace et al. (2007). The lateral domains with different rheological properties are enlightened with different colours. The final scheme comprises: a weak zone under the Elbe Fault System domain (EFS) after the results from PACE (Palaeozoic Amalgamation of the Central Europe) seismic experiment, a domain characterised by a normal Variscan-type crust, a transition domain with a high density lower crust and a stiff East European Craton (EEC) re­gion. Abbreviations: EEC=East European Craton; EFS=Elbe Fault System

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Figure 3.1.21. Direction of the largest horizontal stress for the Central European Basin System as derived from the most recent release in 2005 of the World Stress Map Project (WSMP), Reinecker et al. 2005

cording to the study of Cacace et al. (2007), strong strain localisation occurs without modelling pre-defined faults and/or shear zones. Asymmetric modes of deformation as well as deep penetrating low-angle faults are not needed to explain and reproduce the broad lithosphere deformation style. On the contrary, the heterogeneous pattern of continental deformation may be regarded as the natural consequence of the asymmetric nature of thermal and structural heterogeneities within the lithosphere, (see Fig. 3.1.23). Lateral rheological in-homogeneities have been proven to induce variations in the azimuth of the strain rate eigenvectors as already suggested by previous numerical studies (e.g., Marotta 2005). On the other hand, lateral contrasts turn out to affect the strain-rate eigenvalues revealing their key role in stiffening the propagation of tectonic deformation. Based on a weak lower crust below the Elbe Fault System (panel (c) of Fig. 3.1.23) the results reproduced suggest a rather undeformable area in the northernmost part of the North German Basin in agreement with previous modelling results (e.g., Marotta et al. 2002) and with the local observed a-seismicity of the area (e.g., Cloetingh et al. 2006). In contrast with previous studies (e.g., Gölke and Coblentz 1996), the results also demonstrated that the direction of the principal stress axes is not totally independent of the rheology of the lithosphere. The presence of different structural domains at crustal and shallow mantle level was found responsible for the present-day local variations observed in the direction of the regional stress field, (see Fig. 3.1.24). The local bending in the direction of the maximum horizontal stress components, from a regionally consistent NW-SE orien-

Figure 3.1.22. Triangular horizontal strain-rate eigenvectors derived from the ITRF2000 solutions in Central Europe (blue colours indicate extension, and red colours indicate compression). Light blue and red colours in the left angle indicate non-significant strain rates. Modified after Marotta (2005)

Driving mechanisms for basin formation and evolution

Figure 3.1.23. Horizontal deformation field (i.e. strain rates eigenvectors and eigenvalues) from the numerical study of Cacace et al. (2007). (a) Effects of the sediments and crustal thickness variations (panel (a) and (b) of Fig. 3.1.19) on the style of horizontal deformation. The general trend as geodetically observed (Fig. 3.1.22), is well reproduced. (b) Effects of lateral variations in the lithosphere-asthenosphere boundary (panel (c) of Fig. 3.1.19). Due to the topography of the isothermal boundary, a region of localised strain develops. (c) Effects of lateral rheological heterogeneities on the style of deformation. As a consequence of having considered a second weak crustal zone under the EFS (Fig. 3.1.20),the strongest strain localisation is presented below this area. A second and minor domain of localised compressional deformation continues to characterise the transition zone between the Baltic domain and the Variscan domain

tation throughout almost all northern Europe to a N-S to NE-SW-directed orientation under north-eastern Germany and in Poland, was reproduced by only considering the interplay of lateral variations in both the thermal field and in the thickness of the different structural layers together with lateral rheological contrasts within the lithosphere. All the models described above were restricted to the recent, i.e., present day, tectonic setting. The main rea-

son for focusing on the present day tectonic frame is the availability of several high-quality data sets that can be used to constrain model predictions. However, having once obtained the best estimates of the thermo-rheological structure of the lithosphere, including inherited crustal and lithosperic structures, and having constrained their control on the present day deformation patterns and stress field, it is possible to model basin evolution through time, coupling all the obtained information together with vary-

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Figure 3.1.24. Direction of SHmax and related stress regime (colours) from the study of Cacace et al. (2007). Different colours represent the stress regime calculated by comparing the three components of principal stresses, i.e. σ1, σ2, σ3. Green colours indicate thrust faulting, blue colours strike-slip faulting, and purple colours normal faulting

ing stress boundary conditions derived from plate tectonic reconstructions. In this context the thin-sheet approach, requiring a direct relationship between stress and strain rate by means of material parameters, provides an effective tool.

As an example, the interactions between rheology and stress and their effects on the deformation pattern and basins development were considered by the use of a large scale finite element spherical thin-sheet model for the Mesozoic tectonic evolution of the Central European Basin System. Taking the rheological structure of the lithosphere as derived from the integral study by Cacace et al. (2007), the evolution of major subsidence centres and adjacent areas of uplift through time was reconstructed in relation to small changes at the effective stress boundary conditions as defined by previous plate tectonic studies (e.g., Ziegler 1990; Ziegler et al. 1998; Cloetingh et al. 2006; Scheck-Wenderoth in press). Although complex, the tectonic evolution of the CEBS may be schematically summarised as follows: basin formations during Late Carboniferous-Early Permian times was accompanied by intense magmatism and was followed by a period of thermal subsidence during latest Late Palaeozoic and earliest Triassic. Rifting activity was mainly localised along the East Greenland-Norwegian margins and in the Tethyan domains. It started during the Triassic and led to the establishment of a regional extensional regime throughout the Middle and Late Triassic as well as during the Early Jurassic. Rifting resulted in increased thicknesses of Triassic deposits both in the N-S trending grabens, Central Graben, Horn Graben and Glückstadt Graben, and in the NW-SE striking main Permian basins, North German Basin, Polish Trough and Danish Basin (e.g., Michelsen 1982; Dadlez et al. 1995, 1998b; Clausen and Pedersen

Figure 3.1.25. (a) Imposed boundary conditions in terms of horizontal deformation velocity (yellow arrows) in order to resemble the tectonic setting observed during the Middle to Late Triassic evolution of the Central European Basin System. The eastern boundary is kept free. (b) Vertical deformation pattern (red colours stand for uplift and blue colours for subsidence). The thick red isolines indicate no vertical displacement. Almost the entire basin system undergoes regional subsidence. More local features shown are: (ii) the occurrence of a depocentre of subsidence located under the Polish Trough (e.g. Scheck-Wenderoth and Lamarche 2005) and (ii) an asymmetric trend of vertical deformation with southern areas (the so-called “Southern Permian basin”) showing more volume of vertical displacement with respect to the northern domain (“Northern Permian Basin”)

Driving mechanisms for basin formation and evolution

Figure 3.1.26. (a) Imposed boundary conditions in terms of horizontal deformation velocity (yellow arrows) in order to resemble the tectonic setting observed during the Middle Jurassic evolutionary phase of the Central European Basin System. (b) Vertical deformation pattern (red colours stand for uplift and blue colours for subsidence). The thick red isolines indicate no vertical displacement. Two broad areas which differ in terms of vertical deformation evolve. Uplift dominates within the North Sea domain, in southern Denmark and in the northern Germany, whereas subsidence still prevailed within the other parts of the basin system

Figure 3.1.27. (a) Imposed boundary conditions to resemble the geometry of Alpine dynamics (red arrows) and sea-floor spreading across the Middle Atlantic Ridge (light blue arrows) as observed during Late Cretaceous-Early Tertiary times. (b) Resulting vertical deformation pattern. Red colours stand for uplift and blue colours stand for subsidence. Thick red and blue isolines indicate no vertical displacement

1999; Maystrenko et al. 2005b; Scheck-Wenderoth and Lamarche 2005). Figure 3.1.25 illustrates the boundary conditions imposed in order to resemble the observed tectonic frame, panel a, as well as the resulting modelled vertical deformation patterns, panel b.

Triassic-Early Jurassic regional subsidence was interrupted during Middle Jurassic by a thermal dome centred at the triple junction in the Central North Sea Rift System (i.e., Viking Graben, Moray Firth Basin and Central Graben) which resulted in regional uplift in the Central North Sea domain during Middle Jurassic (“Mid-Cimmerian” unconformity). Middle Jurassic uplift was not lim-

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ited to the North Sea area but also extended across Denmark to northen Germany including the Mid North Sea Ringkøbing-Fyn High and the southern part of the Danish Basin (e.g., Vejbæk 1997) and the northern portion of the North German Basin. After the relaxation of the thermal dome, regional subsidence prevailed in the Late Jurassic and was followed by a period of tectonic silence during Early Cretaceous. Figure 3.1.26 shows a schematic representation of the imposed stress boundary conditions, panel a, and of the geometry of resulting deformation pattern, panel b. By the beginning of the Late Cretaceous, i.e., Turonian, the inception of the Alpine orogeny drove long-wavelength intra-plate compressive stresses throughout whole Central Europe. As a consequence the major basins within the CEBS underwent a phase of inversion while minor rifting activity and related subsidence occurred in the North Sea Domain from Early Oligocene times with a depocentre in the North Sea Rift. Figure 3.1.27 gives a representation of the imposed boundary conditions, panel a, together with the associated deformation pattern, panel b.

3.1.4 Modelling complex basins The Central European Basin System is an exceptional example of a complex sedimentary basin and the application of modelling techniques, as almost all currently available basin models have been applied to it. It is complex in terms of having suffered stretching and inversion phases, being located above a highly heterogeneous crust and even mantle. Last but not least the upper crust tectonics is complex due to the repeated activation of salt tectonics. The models applied account for special aspects of historical events; however, none of them can capture the complex processes causing the evolution of sub-basins and localised inversion zones in detail. The models considered here, of course, are based on basic physical principles which try to reduce the complex reality to a simple basic set of rules in order to understand the underlying mechanisms. In this sense they can be applied to complex basins with some success: they should, however, not be overemphasised concerning details relevant for applications like an oil play. This becomes immediately clear from the example of McKenzie’s one-dimensional stretching model. Although it may be adapted to any given well, it will fail to account for local salt

tectonics, and the linear thermal model contradicts observations and laboratory results: the linear temperature gradient (~30 °C/km) would lead to about 3,000 °C at the transition between the litho- and asthenosphere at 100 km depth, or if the temperature at this transition is about 1,300 °C, as is usually considered, the thermal gradient becomes 13 °C/km. The first value is certainly too high whereas the second is much too low. Similarly, the simple shear model of Wernicke provides only a reasonable concept; however, there are no hard data to support it. The use of buckling models with regard to the continental crust, originally developed for the oceanic crust are limited due to the complex rheology of the continental crust which can support elastic behavior only to a minor degree. Nevertheless, these simple models help to understand specific phases of basin evolution, even when they are complex. The more advanced models such as the thin sheet model, which we may call 2.5 dimensional, provide insight into the interaction between far field boundary stresses or strains and local variations in material parameters, independent of the basic considerations elastic, plasto-elastic or visco-elastic. However, there are unresolved geological problems concerning the evolution and role of major geological discontinuities like faults and/or variations in crustal composition reflecting the unsolved question: “What came first, the chicken or the egg”. More precisely, due to restricted computer capabilities, it is frequently necessary to reduce the 3-D problem to two dimensions by integrating over the lithosphere. A critical point to consider in this is the possible existence of more than one decoupling horizon within the lithosphere, salt always being dominant in this respect. There are no computer codes available which can deal with a basin with an aspect ratio of 1:10 for lithosphere thickness to basin width in sufficient detail. Plus, we must also consider just how detailed our geological and physical data are, particularly those describing the deeper crust and the mantle. Again, even these advanced models only allow us to describe how an inhomogeneous lithosphere will guide the strain vectors and may cause strain localisation which will finally lead to failure at predetermined locations. Even if in the future we are able to employ detailed three-dimensional models, they will be limited by our restricted knowledge of rock properties and their real distribution. In conclusion, we believe that the models we have presented provide useful tools to aid in understanding past and present processes. It is, however, imperative not to accept the modelling results at face value but to realise they are limited by the assumptions made concerning the history of the basins and by inadequate geological data.

3

Chapter 3.2

Crustal structures and properties in the Central European Basin system from geophysical evidence C.M. Krawczyk · W. Rabbel · S. Willert · F. Hese · H.-J. Götze · D. Gajewski & the SPP-Geophysics Group*

3.2.1 Introduction This subchapter will present selected results from different geophysical observations and studies with the aim (i) to show the potential of each geophysical method alone, (ii) to outline how the different methods can supplement each other, and (iii) finally, to combine the different attempts into one integrated geophysical model, linked ideally to geologically derived models. Therefore, first of all, we present structural images, secondly, these are interpreted to yield a geometric framework for any kind of subsurface model, and thirdly, if the data allow, physical parameters will be derived. At the end, this will all serve along with other constraints as input for modelling and simulation studies.

To work on this concept, we mainly focus here on results from studies that are ideally suited to address the aforementioned aims and the structures controlling the Central European Basin system (CEBS). We are not deliberately omitting all the other extremely valuable investigations available in the area (cf. this book and references cited therein). One of those structures where a series of different geophysical investigations is available is the crustal-scale Glückstadt Graben (GG) which serves as a recurrent theme throughout this chapter. The second recurrent theme discusses the Avalonia-Baltica Suture. This crustal and even lithospheric-scale feature has also been investigated by different geophysical methods in order to reveal its deeply buried trace. The chapter itself is built around the different geophysical

seismology -4

-2

Frequency [Hz]

10 - 10

10-2 – 10-1

10-1 - 100

Wavelength [km]

26.000-800

800-80

80-8

Source

oscillation of the earth

long-distance earthquakes

short-distance earthquakes

Application

entire earth

Depth range [km]

104

103 - 102

102 - 101

Resolution [km]

102

101

101 - 100

deep

mantle/crust

exploration

seismics

shallow seismics

laboratory

Frequency [Hz]

1 – 10

10 – 100

10 - 100

101 – 103

102 – 104

104 - 106

Wavelength [km]

6 – 0.6

0.3 – 0.03

0.1 – 0.01

10-1 – 10-4

10-2 – 10-4

10-4 – 10-6

Source

earthquakes explosions

Application

upper/crust

upper crust

short offset

engineering

mining

models

Depth range [km]

102 – 101

101 – 100

100 – 10-1

102 – 101

101 – 100

100 – 10-1

Resolution [km]

101 – 100

10-1 – 10-2

10-2

101 – 100

10-1 – 10-2

10-2

seismic shots and non-explosive sources

non-explosive sources

non-/explosive and ultrasonic sources

Figure 3.2.1. Overview of different seismic methods and their scales

*(includes also: M. Baykulov, F. Bilgili, H.-J. Brink, H. Busche, L. Hengesbach, N. Hoffmann, C. Hübscher, H. Jödicke, Y. Maystrenko, P. Schikowsky, S. Schmidt, F. Theilen, T. Yegorova, M. Yoon, H. Zöllner)

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methods with their different structural resolution and technical background.

3.2.2 Structural inventory and physical properties from seismic observations 3.2.2.1 Overview Seismic methods are based on the propagation, reflection and refraction of elastic waves in the subsurface. Thus, the seismic images we produce in the first instance show the heterogeneities in the elastic properties of the traversed volume. These differences in heterogeneity may be imaged as reflectors, refractors and converting interfaces, and are attributed in the subsequent interpretation to specific structures and boundaries. Since all this is based physically on the wave equation, any seismic working step must be considered frequency dependent. This is important as frequency and wavelength respectively define both the vertical and horizontal resolution of the selected seismic method. Different seismic methods with typical wavelengths and applications are summarised in Fig. 3.2.1. A complete seismic project generally encompasses acquisition of field data, data processing, particularly seismic imaging, and seismic interpretation. During field data acquisition a signal is emitted at a certain time, and transducers deployed spatially or along a profile line record the signals that are received back. The measured physical parameter is the travel-time of the elastic wave through the medium, which needs to be translated into depth during data processing via a velocity model. Velocity-model building and the migration of the data are still, as they have been for decades, the most sensitive and critical steps in seismic processing, in order to derive a true, or at least best possible, constrained earth model. During data interpretation, well data or other constraints are essential for horizon calibration. Reflection seismic data also make it possible to characterise the picked horizons and fault surfaces with respect to seismic attributes like amplitude, coherence, dip, azimuth and curvature. Specialised and cutting edge techniques in acquisition and processing are multi-component techniques (e.g., Stewart et al. 2002; Barkved et al. 2005; Paulus et al. 2005; Maxwell and Criss 2006), full wave form inversion (Pratt 1999), wave-equation migration (e.g., Biondi 2006), optimised aperture migration (Schleicher et al. 1997; Lüth et al. 2005), time-lapse imaging (Ebrom et al. 1998; McKenna et al. 2001; Yuh 2003 and references therein), multi-parameter stacking with its derivatives

and extensions (e.g., Box 3.2.1-1), and receiver functions techniques (Box 3.2.1-2). Besides these, a vast amount of comprehensive articles and text books is available, of which we recommend the following for further reading: Telford et al. (1976); Meissner and Stegena (1977); Claerbout (1985); Militzer and Weber (1985); Sheriff and Geldart (1995); Stein and Wysession (2003); Aki and Richards (1980); Yilmaz and Doherty (2001); Robein (2003); Biondi (2006); Brown (2004); Chopra and Marfurt (2006).

3.2.2.2 Detailed structural architecture and tectonic history from reflection seismics The reflection seismic method is the geophysical technique most important for structural imaging within the Earth, applicable in three dimensions, from m- to km-scale, and on land as well as at sea (cf. Fig. 3.2.1). Pulses of seismic energy are emitted and signals are reflected back to the surface at structures and interfaces in the subsurface, depending on impedance contrasts met. The recorded signals need extensive processing before a seismic section is produced that may be used for further interpretation. Limiting factors are the knowledge of the velocity model of the studied volume, side-reflections, signal-to-noise ratio, and steep dips. However, since reflection seismic sections allow geological models to be built, the main applications are exploration purposes, shallow geotechnical tasks, and research. Other important branches are seismic stratigraphy, used mainly in sedimentology, seismic attribute determination (both also addressed in this book), and the derivation of physical parameters like elastic constants or porosity. A variety of atlases of the NW and Central European region are available. Seismic data were published by DEKORP (Meissner and Bortfeld 1990) and by BIRPS (Klemperer and Hobbs 1991; Snyder and Hobbs 1999; Brewer et al. 1983) as well as in a range of individual papers (see below). Krawczyk et al. (2008) provide a comprehensive review of geophysical results in the light of Caledonian tectonics in Central Europe, thus also covering selected parts of the Southern Permian Basin. Since all this information is available, and new results from shallow reflection seismics are also presented elsewhere in this book, we will concentrate here on major results from sites covering the basin between the North Sea and the Baltic Sea, and from crustal levels down to the uppermost mantle (all sections are also found on the CD attached to this volume). In the North Sea, a vast number of 2D and 3D reflection seismic surveys are available, some from research

Crustal structures and properties in the Central European Basin system from geophysical evidence

Box 3.2.1-1 Specialised seismic processing and acquisition concepts - 1 The Common Reflection Surface (CRS) stack The Common Reflection Surface (CRS) stack is a multi-parameter stacking technique that requires no a priori macro-model input. The traveltime t of a reflection event is described by the three parameters A, B, and C (Eq. 3.2.1), representing subsurface location, curvature and dip of a reflector. In contrast, the classical Common Mid Point (CMP) stack uses only one parameter, the move-out velocity Vnmo (Eq. 3.2.2) assumed to be equal to the stacking velocity. With X - offset, m - midpoint coordinate, and t0 - zero-offset two-way traveltime, the stacking formulae are simplified to: CRS-Stack:



(3.2.1)

CMP-stack:



(3.2.2)

Both equations represent short-spread hyperbolic traveltime approximations applicable to any kind of heterogeneous model. Therefore, X should be reasonably small to satisfy the hyperbolic assumption. Due to this limited set of assumptions, stacking is a stable procedure in seismic data processing. The CMP stack is a subset of the more general CRS stack. If the midpoint coordinate m is zero we immediately transform equation 3.2.1 into equation 3.2.2 with Vnmo=1/t0C. The advantages of the CRS stack over the CMP stack are given by (i) the incorporation of the midpoint coordinate m and (ii) the considered dip of the reflectors when building super bins. This leads to a considerably larger number of traces in the CRS stack and, thus, to an improved signal to noise ratio and to a higher reflector continuity. Practically, both stacks follow similar procedures. Prior to the stack the stacking parameters need to be determined. In the CMP stack we determine Vnmo by testing a certain range for this parameter and choose its optimum value by evaluating a coherence measure like the semblance. This process is usually described as velocity analysis. A similar approach is applied in the CRS stack; however, we have to fit three parameters simultaneously. The processing parameters Vnmo and A, B, C can also be translated to physical parameters. The CMP parameter move-out velocity can be transformed into an interval velocity via Dix inversion in a horizontally stratified medium. The CRS parameters A, B, and C are related to wave-front curvatures. These are important for advanced processing steps like spreading corrections, reflection tomography or depth imaging.

Further reading Initial studies on multi-parameter stacking were performed  by Gelchinsky (1988), and altered to the CRS method a decade later by Hubral (1999) and the Karlsruhe group: the most detailed descriptions with implementation details and many synthetic and real data examples are found in Müller (1999) and Mann (2002), who also show extensions of the CRS technology. Duveneck (2004) applied CRS parameters to velocity model building by reflection tomography.

investigations but mainly from commercial applications, since it is a region with a high density of exploration targets. Since oil reservoirs can only be expected within the sediments and costs for drilling oil increase significantly with depth, the industry’s main focus lies in the first few thousand meters of consolidated sediments. Consequently, data acquisition and processing are optimised for imaging the economically relevant part of the subsurface. Scientific targets, however, also focus on other aspects and include near-surface geophysics (e.g., for engineering purposes, cf. Wilken et al. 2008) or go much deeper and include the whole crust and the upper mantle (e.g., for tectonic studies, see below). Here, we show two examples of industry-acquired marine reflection seismic data from TGS NOPEC that accu-

rately image the basin sediments underneath the German North Sea sector in great detail with 10-20 m horizontal and vertical resolution (see Fig. 3.2.2 for location). The spectacular sequence and shape of the geological units, which are visible in these time-migrated seismic sections (Figs. 3.2.3, 3.2.4), formed over more than 300 million years by the continuous influence of different, sometimes simultaneous geological processes such as sedimentation, subsidence, salt tectonics, and regional tectonics, leaving behind remnants of structures which allow the reconstruction of those processes. Therefore, seismic sections not only provide in-situ information on the deeper subsurface but, by interpretation and retro-deformation of strata and structural remnants, also give an insight into the variable geological and tectonic history of the region (see Kley et al. this volume for examples).

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Box 3.2.1-2 Specialised seismic processing and acquisition concepts - 2 Receiver Functions A receiver function is the seismic response function of the subsurface beneath a seismograph with respect to an impinging body wave emitted from a remote source. It represents an amplitude-time function usually computed from three-component seismograms of long-distance earthquakes (typically 30°-90° epicentral distance). So-called P-receiver functions are widely used for imaging the crustal structure down to the Moho level as well as for investigating the upper mantle discontinuities. P-receiver functions are understood as P-to-S converted waves which are generated at interfaces in the subsurface at the moment when the up-diving P-wave crosses them (Fig. a). Since S-waves travel with lower velocity than P-waves, the converted S-waves arrive at the receivers with a travel-time delay depending on the depth of the interface and the seismic velocities of the hanging wall (Fig. b). For a horizontal interface at depth z the travel time delay ΔtSP of converted “S minus P” reads tSP = z·[(vs-2-p2)1/2-(vP-2-p2)1/2] where vP and vS are the P- and S-wave velocities of the hanging wall and p is the slowness of the incoming waves (i.e., the slope of the travel-time curve).

P- and S-waves are polarised nearly orthogonally to each other so that they can be separated by computationally adjusting the orientation of the receiver coordinate system. The P-receiver function is obtained by a deconvolution process which removes the earthquake source signal, represented by the P-wave form, from the converted S-wave signal. Alternatively, a cross-correlation function of the P- and converted S-wave signal could be computed. This results in amplitude-time functions where the P- and S-wave signals appear, ideally, as spikes at time 0 and at the “S minus P” travel time delay (Fig. c). Finally, by assuming an adequate S-wave velocity-depth function, the amplitude-time functions can be converted to a depth-image of the interface by applying a migration process which accounts for the correct ray path of the incoming wave (Fig. d). For a horizontal interface at depth z the conversion point or image point is found at a horizontal distance ΔxS from the receiver position given by ΔxS = z·p/(vs-2-p2)1/2. Because converted signals are rather weak receiver functions of many earthquake arrivals have to be summed (“stacked”) in order to improve the signal to noise ratio. Stacking can be performed after or before migration. In the latter case a time correction has to be applied to the records before stacking in order to consider the different incidence angles of the waves to be stacked.

Further reading Langston (1977); Owens et al. (1984); Ryberg and Weber (2000); Vinnik (1977) and Yuan et al. (1997).

The seismic section shown in Fig. 3.2.3 comprises several seismostratigraphic features within the German North Sea and images the two large grabens in this area. The Central Graben developed in the Triassic shortly after the creation of the Horn Graben, which already existed in the Late Permian. These grabens are the re-

sult of extensional tectonics, which was caused by the break-up of the Pangea supercontinent (e.g., Ziegler PA 1988). In the grabens, Permian evaporites with an assumed initial thickness of about 2500 m are overlain by Triassic sediments, which give, after interpreting changes in sedimentation thickness and onlaps at the flanks of

Crustal structures and properties in the Central European Basin system from geophysical evidence

the grabens, the information needed to date the break-up process of the Central Graben.

Figure 3.2.2. Location map of three selected reflection seismic lines in the German North Sea sector. Profiles g28 and g24 are industry lines acquired by TGS NOPEC. Recording of 7 s allowed imaging of crustal structures down to the pre-Permian. The scientific, deep seismic MONA LISA 1 profile with a record length of 25 s reached the mantle

The first mobilisation of the Zechstein salt during the extension process led to the creation of salt diapirs and salt walls that exerted structural control on subsequent sedimentation in the area. Salt diapirs formed predominantely in close proximity to the grabens. Here, salt intruded into the spaces created by regional extension (Fig. 3.2.3). However, salt diapirs did not originate close to the West-Schleswig-Block. Here, listric faults formed but salt movement was not triggered (Fig. 3.2.4). Possible reasons for this are the initial thickness of salt, which is remarkably high in the grabens, the depth and thermal differences between both locations. Common features within the subsurface under the German North Sea sector are “touchdown-structures” (Fig. 3.2.4), where younger sediments directly overlie the pre-salt sediments while salt is missing. Some of those structures seem to be related to basement faults, some seem to be a result of salt movement and depletion.

Figure 3.2.3. Time-migrated reflection seismic profile g28 across the German sector of the North Sea running from the Pen Handle region towards the island of Heligoland (data from TGS NOPEC). Prominent seismostratigraphic features of the southern North Sea are highlighted (see text for details)

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In the Early Jurassic, a rise in sea level allowed shallow marine sedimentation. The Middle Jurassic to Early Cretaceous uplift, by compression or transpression (Kockel 2002; Ziegler 1990), led to the erosion of Late Triassic and Early Jurassic sediments. The uplift was succeeded by sedimentation of clastic (Early Cretaceous) and marine sediments (Late Cretaceous until Palaeogene). Erosional truncations are caused by Cretaceous transgression (Figs. 3.2.3, 3.2.4). In the Late Cenozoic the simultaneous uplift of the Fennoscandian Shield and accelerated subsidence in the North Sea Basin led to enormous deposits of fluvial sediments (e.g., Fig. 3.2.3), with a volume of about 62x103 km3 over a surface of 28x103 km2 in the southern North Sea Basin (Overeem et al. 2001). The correlation of geochemical records and seismic analysis reveals that the average rate of sediment supply and the channel characteristics were controlled by climate (Overeem et al. 2001). Where technically and economically possible, seismic information is combined with that derived from drilling samples that facilitate dating. But in places it may be too costly and even technically impossible to drill deep enough. However, seismostratigraphic correlation is often possible. Over large areas in the Southern Permian Basin and the North Sea it is especially feasible, since most of the seismic horizons have an extremely continuous and stable reflectivity pattern that can be traced. In addition, seismic interpretations can be combined with other information from the local geology or other geophysical methods where available (see below).

The deeper crustal and upper mantle levels of the North Sea were investigated, among others, by the MONA LISA survey (MONA LISA Working Group 1997a,b). Of course, lateral and vertical resolution of the data are decreased compared to the reflections shown above, but spectacular deep seismic reflections from the upper mantle were observed along the MONA LISA profiles in the southern North Sea (Fig. 3.2.5, see also Fig. 3.2.19 below). These clearly visible reflections recorded discontinuities down to 80 km depth. The coherent reflection elements form patterns suggesting a continuously southdipping interface. The origin of these mantle reflections is still unknown despite many analyses and modelling efforts. Current hypotheses for the interpretation of the deep reflections, which are based mainly on seismic amplitude analysis and tectonic plausibility, include (i) a fossil plate interface remnant of the Caledonian continent collision connected with the subduction of the Tornquist Ocean, (ii) shear zones formed during Late Palaeozoic extension possibly connected with fluid-filled fracture planes, and (iii) seismic anisotropy caused by lattice preferred orientation of mantle minerals in a tectonic creep process (see Abramovitz et al. 1998; Abramovitz and Thybo 2000; Balling 2000; Krawczyk et al. 2008 for further discussion). The Glückstadt Graben (GG), located further eastwards in the CEBS, is another important feature that was analysed seismically in more detail for both exploration purposes and under crustal-scale geodynamic considera-

Figure 3.2.4. Time-migrated reflection seismic profile g24 in the North Sea (data from TGS NOPEC). Highlighted structural features are discussed in the text

Crustal structures and properties in the Central European Basin system from geophysical evidence

Figure 3.2.5. Deep seismic reflection profile MONA LISA 1 (time-migrated) showing the lithospheric structure of the southern North Sea down to 25 s TWT corresponding to ca. 80 km depth (modified after Abramovitz et al. 1998 and Domaschk 2000). Note the strong reflectivity in the upper mantle. RFH – Ringkøbing-Fyn High, HG – Horn Graben, WSP – West-Schleswig Platform, NDB – North German Basin

tions (for location see Fig. 3.2.6). The GG is situated between the Caledonian Deformation Front in the north and the Elbe Line in the south. It suffered the main tectonic stages of the Central European Basin System (Maystrenko et al. this volume; Littke et al. this volume) with some superposed local features. The GG area was a centre of subsidence at the end of the Permian, and the initially thick Permian salt moved into salt walls, diapirs and pillows during post-Permian times. The GG as it appears today

was initiated in Triassic times with major subsidence and deposition of up to 9 km thick sediments in its central part (Maystrenko et al. 2005a). Late Cretaceous-Early Tertiary compression caused increased salt activity (Kockel 2002; Maystrenko et al. 2006). Available data sets across the GG were acquired and processed in the 80s with the main focus on the sedimentary fill of the basin. Here, we discuss selected seismic

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C.M. Krawczyk · W. Rabbel · S. Willert · F. Hese · H.-J. Götze · D. Gajewski & the SPP-Geophysics Group Figure 3.2.6. Schematic tectonic map of the Glückstadt Graben area with the location of the reflection seismic profile GG1 shown in Fig. 3.2.7 (after Yoon et al. 2007)

lines reprocessed using the Common Reflection Surface (CRS) stack method (see Box 3.2.1-1 for detail). This application first concentrates on the lower crustal structure in order to investigate the influence of old, deep rooted processes on the evolution of the Central European Basin System (CEBS), and secondly it emphasises the imaging of salt plug families and the sedimentary structure. A comparison between CRS stack and the more conventional CMP stack, as seen in the example of profile GG1 shows that the CRS stack generally improved the image quality of the sedimentary cover as well as that of the internal salt structures (Fig. 3.2.7). Also, the visibility of reflections from the crystalline crust and from the Moho was significantly increased (Yoon et al. 2007). Thereby, the CRS stack of profile GG1 indicates a more or less flat Moho topography along the GG. Even though a stacked section cannot be used to derive true dips and angles with-

out migrating it, a more or less flat lying Moho would contradict former interpretations which proposed a Moho uplift of up to ca. 4 km underneath the graben axis in the area, thus suggesting an extension model (Bachmann and Grosse 1989; Brink et al. 1990). Further support for such a flat Moho could be cited by the flat Moho already observed on the depth-migrated BASIN9601 profile crossing the NE German Basin (DEKORP-BASIN Research Group 1999; Krawczyk et al. 1999). Since this finding is in conflict with the former extension model of the GG and other areas, the geodynamic model of the region may have to be revised. In addition, one could also speculate that the present-day flat lying Moho could also represent a different phase of petrophysical equilibrium that does not correspond to the Moho during basin initiation. This would mean that the syn-rift uplifted Moho flattened during post-rift thermal subsidence.

Crustal structures and properties in the Central European Basin system from geophysical evidence

Figure 3.2.7. Reflection seismic profile GG1 across the Glückstadt Graben (for location see Fig. 3.2.6; after Yoon et al. 2007). (a) CMP stack and (b) CRS stack, each with inset showing zoom of Moho reflections. Two zooms of salt reflections are given in (c), with left column - CMP stack, right column - CRS stack

With respect to the upper crustal imaging of salt plug families and the sedimentary structure in the GG, new migrated images provide structural details which may also suggest an alternative and more speculative view on the structural setting of the area: The Jurassic salt plug indicates tectonics similar to observations in the Allertal region east of the GG, where overthrusting plays a major role in the evolution of the salt structures (cf. Kley et al. this volume; Lohr et al. 2007a). As a consequence, the classical concept of halokinesis as applied to Northern Germany (Trusheim 1957; Sannemann 1963; Jaritz 1973) also has to be altered at least for the early stages (i.e., Early Triassic, Buntsandstein). In the GG, salt structures and overthrusts are tectonically associated. Both features developed predominantly during the inversion of the Lower Saxony Basin and define its northern fringe. The kinematic evolution of the GG is also partly discussed in Kley et al. (this volume). To shed further light on the nature and origin of the interpreted features made apparent by the different processing attempts, coincident magnetotelluric observations and gravity modelling can be considered (see below). Across the NE German Basin, the first complete structural image of both Variscan- and Caledonian-influenced

areas of the North German Caledonides was provided by the BASIN’96 survey (DEKORP-BASIN Research Group 1999; Krawczyk et al. 1999), calibrated in the Mesozoic part by numerous wells (Hoth et al. 1993b). The survey revealed that the Moho is a continuous structure across the entire BASIN9601 profile (see part of it in Fig. 3.2.8) and is located at the base of a ca. 2 – 4-km-thick reflector band at 30 km depth beneath the entire basin (35 km depth at the margins). The Variscan structures along the southern basin margin show the basal Zechstein reflector with ca. 5 km vertical offset at the Gardelegen Fault (Kossow et al. 2000). Offsets of 2.5 km are observed at the Haldensleben Fault, and at the Harz Boundary Fault, underlain by strong mid-crustal reflectors at c. 20 km depth. These offsets were geometrically modelled to mark the mid-crustal expression of the fault zones observed at surface (Kossow and Krawczyk 2002). At the northern margin of the basin, the Zechstein reflectors shallow from 3 to 1.5 km depth and are truncated by the overlying Triassic. The landward extension to the N of profile BASIN9601 across Rügen, however, could not resolve structures below the Mesozoic basin (Mayer et al. 2000). Despite this, moderate-amplitude SW-dipping reflectors are observed down to ~20 –25 km depth on BASIN9601, beneath the northern third of the basin (Fig. 3.2.8), and this surface is interpreted as the Caledonian suture. The wedging form of this structure

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Figure 3.2.8. Composite from migrated reflection seismic profiles BASIN9601, PQ2-9.1 and PQ2-5 (top) with interpreted linedrawing (middle) and refraction seismic velocity model (bottom) showing the extent of Baltica between the SW Baltic Sea and NE German Basin (combined after DEKORP-BASIN Research Group 1999 and Krawczyk et al. 2008). G14 labels a well in the Baltic Sea

Crustal structures and properties in the Central European Basin system from geophysical evidence Figure 3.2.9. a) Cenozoic isochore map summarising the interpretation of the entire NeoBaltic seismic grid; the location of the example profile given in b) is marked. b) Time-migrated seismic section HE02-33 and interpretation in the western Baltic Sea area, southwestern Bay of Kiel (after Hansen et al. 2005). Immediately above the salt dome, faults cut from Top Zechstein up to the surface, minor faults are observed in the Cenozoic. This reveals the thin-skinned tectonic response of the Post-Permian supra-salt succession to the major plate tectonic periods. BMT – Base Middle Triassic, BUT – Base Late Triassic, IUT1, 2 – Internal Late Triassic layers, MJU – Mid Jurassic Unconformity, BUC – Base Late Cretaceous, IUC1, 2 – Internal Late Cretaceous layers, BC – Base Cenozoic, IC1, 2 – Internal Cenozoic layers. Arrows mark toplap terminations against erosional unconformity (after Hansen et al. 2005)

suggests that Baltica crust extended much farther southward below northern Germany (i.e., into Avalonia) than previously thought, possibly as far S as the basin centre or the Elbe Line (Krawczyk et al. 1999). The Baltic Sea hosts the NE boundary of the Central European Basin with a much smoother shallowing of Permian strata on top of the Palaeozoic crust than observed generally at the southern basin margin. Several regional reflection seismic surveys in the SW Baltic Sea have focused on the uppermost representation of sedimentary sequences in order to map the Mesozoic evolution of the area (e.g., Erlström et al. 1997 and older references therein) or to investigate the hydrocarbon prospectivity (e.g., Petrobaltic survey, Rempel 1992; JEBCO survey, unpubl.), or both (SASO project, Schlüter et al. 1997; Piske et al. 1994).

More recently, the NeoBaltic offshore surveys were acquired (Hübscher et al. 2004). These high-resolution surveys focussed on the evolution of the post-Zechstein successions in the Bay of Kiel area with emphasis on the important neotectonic activity and its relation to salt dynamics and the present-day stress field. The data and derived isochore maps reveal the major sequence boundaries traced across the south-western Baltic Sea (Fig. 3.2.9). The NeoBaltic seismic sections elucidate the thinskinned tectonic response of the Post-Permian supra-salt succession to the four major plate tectonic periods which characterise this realm (Hansen et al. 2005, 2007). During the Triassic and the Early Jurassic, E-W extension and the deposition of clastic sediments initiated the movement of the underlying Zechstein evaporites. The deposition ceased during the Middle Jurassic, when the entire

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area was uplifted as a result of the mid North Sea doming. The uplift resulted in pronounced erosion of Late Triassic and Early Jurassic strata. This event is marked by a clear angular unconformity on all the seismic sections. The region remained an area of non-deposition until the end of the Early Cretaceous, when sedimentation resumed in the area. Throughout the Late Cretaceous, the sedimentation took place under tectonic quiescence. Reactivated salt movement is observed at the Cretaceous-Cenozoic transition as a result of the change from an extensional to compressional regional stress field. The vertical salt movement influenced the Cenozoic sedimentation and resulted in thin-skinned faulting (e.g., Fig. 3.2.9). As part of the Petrobaltic survey, the reflection seismic grid in the Rerik area of the coastal Mecklenburg Bay was reprocessed and interpreted (Zöllner et al. 2007). With 1x2 km grid spacing and two sea-land transition profiles the connection to an onshore profile grid and drill holes is permitted. In addition to mapping stratigraphic horizons and determining the structural evolution of the area over time (see Kley et al. this volume), it was possible to detect three salt pillows, which were previously unmapped in the area (Fig. 3.2.10). These are responsible for the creation of smaller subsidence centres and angular unconformities in the Late Triassic Keuper, especially in the vicinity of the fault-bounded Grimmen High. In this area, partly Early Jurassic sediments overlie the Keuper unconformably. The change

from extension to compression in the regional stress field remobilised the salt, leading to a major unconformity marked at the base of the Late Cretaceous. The Base Zechstein is characterised by a system of NW-SE and E-W striking faults with little vertical displacement but an increasing number of faults towards the east. In the central part of the area the entire Mesozoic sedimentary sequence is cut by a fault. Significant unconformities were found between the middle Keuper horizons and below the base of the Late Cretaceous. Further east, the data from the PQ2 survey provide wellconstrained images of upper crustal reflectivity and lower crustal/uppermost mantle reflections in the SW Baltic Sea (DEKORP-BASIN Research Group 1998; Meissner and Krawczyk 1999; Krawczyk et al. 2002). In the area of the Caledonian suture, a reflection pattern was observed with opposing dips in the upper crust and the uppermost mantle (Fig. 3.2.8). Detailed analysis of dipping reflections in the upper crust has provided evidence of two different sets of reflections, which are separated by the O-horizon, which is the main decollement of the Caledonian Suture Zone (cf. also Piske et al. 1994; Schlüter et al. 1997). S-dipping reflections above the O-horizon are interpreted as Caledonian thrust structures. Beneath the O-horizon, SW-dipping reflections in the upper crust are interpreted as ductile shear zones and crustal deformation features that evolved during the Sveconorwegian Orogeny (see Krawczyk et al. 2002).

Figure 3.2.10. Mecklenburg Bay area in NE Germany with salt structures mapped both onshore and offshore. The offshore continuation of the already known onshore salt pillow Kröpelin is clearly evident and the continuation of the Travemünde pillow suggested at the western margin of the seismic grid. Two new pillows are revealed: the Lake Boltenhagen pillow and further east the Lake Kühlungsborn pillow (after Zöllner et al. 2007)

Crustal structures and properties in the Central European Basin system from geophysical evidence

The Moho reflections at 28-35 km depth appear to truncate a N-dipping mantle structure, which may represent remnant structures from Tornquist Ocean closure or latecollisional compressional shear planes in the upper mantle. A contour map of these mantle reflections indicates a consistent northward dip, which is steepest where there is strong bending of the Caledonian deformation front (Krawczyk et al. 2002). With these new data, a revised geodynamic model was established for the southern Baltic Sea: the collision of Avalonia and Baltica is a thin-skinned, rather than a thickskinned, feature (Krawczyk et al. 2002, 2008; Lassen et al. 2001). The thin-skinned character of the Caledonian deformation and the fact that N-dipping mantle reflections do not truncate the Moho would suggest that the Baltica crust was not mechanically involved in the Caledonian collision and, therefore, escaped deformation in this area. For the sub-Moho reflections, attributed to the traces of Baltica subduction, a mismatch becomes evident across the CEBS. Observations in the North Sea along the MONA LISA lines (see Fig. 3.2.5) imply a southward subduction, while upper mantle reflections in the Baltic Sea along the PQ2 lines would support a north-directed subduction polarity (Fig. 3.2.8). In any event, further wide-angle seismic and receiver functions experiments were conducted as well as gravimetric modelling and MT inversion performed to reveal the possible extent and nature of the underlying Palaeozoic basement (see below).

3.2.2.3 Crustal properties observed by seismic refractions and wide-angle reflections The seismic refraction method makes use of the principle that seismic waves can travel along interfaces in the subsurface that separate layers of different velocity from each other. Applied in either shallow application for soil classification or crustal penetrating set ups (cf. Fig. 3.2.1), the wide-angle seismic method yields velocity profiles of the subsurface that can characterise typically different soil or rock types on a 10 m to km resolution scale, i.e., the Moho is usually recognised as a velocity contrast, from velocities below 8 km/s to values higher than 8 km/s. The refraction seismic velocity models may be further used to depth migrate crustal reflection seismic observations for those parts of the data, where no reflection seismic velocity model building is possible because of insufficient move-out and/ or too short observation offset. Here, we consider the lithosphere scale in order to derive an overall picture of the crustal thickness in the CEBS (see also newly compiled crustal

thickness map for Central Europe by Tesauro et al. 2008), and in order to show the depth expression of the main crustal-penetrating shear zones and terrane boundaries. Besides the rock-type classification itself, the velocity boundaries found from wide-angle measurements are further used as input for gravimetric or geodynamic modelling. Major wide-angle seismic surveys crossing Central Europe were carried out starting with the European Geotraverse project (EGT; key references include works by Blundell et al. 1992 or Dallmeyer et al. 1995), the BABEL and BASIN projects in the Baltic Sea area (BABEL Working Group 1993; DEKORP-BASIN Research Group 1998; Bleibinhaus et al. 1999; Krawczyk et al. 1999) and the comprehensive results of the EUROPROBE Trans-European Suture Zone project (e.g., Pharaoh et al. 1997; Thybo et al. 1999; Winchester et al. 2002; Gee and Stephenson 2006). North of the Elbe Line (EL), the crustal velocity structure changes in the NW and the NE German Basin (Rabbel et al. 1995; Bleibinhaus et al. 1999). The most striking feature is a ca. 10 km thick high-velocity layer in the lower crust between ca. 20 and 30 km depth. Its P-wave velocity is 6.9 – 7.5 km/s, which is typical for shield crust. This high-velocity lower crust extends between EL and the CDF (Rabbel et al. 1995). Other surveys in the North German Basin support this finding (Aichroth et al. 1992; EUGENO-S Working Group 1988; Schulze and Lück 1992). While Hoffmann (1990) interpreted the area S of the EL as an additional microcontinent in the Caledonian collision, its interpretation as an important tectonic element was only later stated more precisely as marking the southernmost extent of Baltica evident in the NE German Basin (DEKORP-BASIN Research Group 1999; Krawczyk et al. 1999; Bayer et al. 1999). Highvelocity lower crust is also found further to the W in the North Sea on MONA LISA profile 1 (Abramovitz et al. 1999; see also Fig. 3.2.19), and to the E in Poland in the POLONAISE97 data (e.g., Grad et al. 1999). In the central part of the North Sea, new results of integrated seismic, gravity, and magnetic studies (Lyngsie et al. 2006) imply that the location of the Central Graben and Viking Graben rift system coincide with the upper crustal suture between Avalonia, Laurentia and Baltica. Lyngsie and Thybo (2007) argued that high-velocity Baltica lower crust extends across the entire northern part of the North Sea, reaching the western edges of the Viking and Central grabens. Oblique thrusting of Avalonia crust over Baltica lower crust corresponds to a 150 km wide zone, outlining the Avalonia-Baltica Suture Zone. The magnetic susceptibility and the density across the Caledonian suture zone range from almost zero and 2715 kg/m3 in Avalonia crust to 0.05 SI and 2775 kg/m3 in Baltica crust. Therefore, the Avalonia-Baltica Suture Zone is characterised by the

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fundamental differences in physical properties between crustal structures of Precambrian origin (Baltica) and Caledonian origin (Avalonia and Laurentia). This inhomogeneous crustal structure attributed to Caledonian collision played a major role in the localisation of the Viking and Central graben-structures during the Mesozoic rifting (Lyngsie and Thybo 2007). The lower crustal structure and pre-Permian evolution of the Glückstadt Graben is known from profile 1 of the EUGENO-S segment of the European Geotraverse project (EUGENO-S Working Group 1988; Blundell et al. 1992). Due to acquisition geometry, a rather schematic model is given for the GG. Crustal velocities increase with depth from ca. 6.4 to 6.7 m/s, and the Moho is found at 28 –30 km depth (Blundell et al. 1992). This is in agreement with results from industry reflection seismic lines (e.g., Dohr et al. 1989). According to various interpretations, the GG rests on lower crust consolidated during the Caledonian Orogeny, lying at the margin of

Baltica (Thybo 2001; Bayer et al. 2002 and references therein). One of the key results from the PQ2 survey in the southern Baltic Sea combined with the onshore BASIN9601 profile is that the extent of Baltica below the NE German mainland is resolved now, both structurally and also in terms of wide-angle velocity modelling (Fig. 3.2.8; Krawczyk et al. 1999; Bayer et al. 1999). A high-velocity lower crust with P-wave values of 7.0 km/s indicates strong Baltica affinities (Bleibinhaus et al. 1999), comparable to the finding along the EGT in the NW German Basin (Rabbel et al. 1995), even though the structural continuity is disturbed by a velocity decrease in an area N of the Caledonian Deformation Front. This coincides with the area of sporadic, not from a single continuously reflecting element generated reflections in the uppermost mantle. These strong reflections apparently dip to the NE, and may mark the traces of Tornquist Ocean subduction (Meissner and Krawczyk 1999). Such reflections are also

Figure 3.2.11. Composite cross section of the Trans-European Suture Zone along the TOR profile (for location see inset). 0-70 km depth: Migrated P-receiver function image of the crust, Moho and uppermost mantle (after Gossler et al. 1999); major tectonic units of the crust are indicated on top of the figure. 70-250 km depth: Tomogram of P-wave velocity based on teleseismic travel time picks of Shomali et al. (2002) (tomographic algorithm see Busche 2001); P-wave velocity depth function of global average lithosphere shown for comparison on the right (IASPEI 91 standard earth model after Kennett 1995). Abbreviations: CDF – Caledonian Deformation Front, EL – Elbe Line, LAB – tentative depth of the lithosphere-asthenosphere boundary (after Arlitt et al. 1999), STZ – Sorgenfrei-Tornquist Zone, VDF – Variscan Deformation Front

Crustal structures and properties in the Central European Basin system from geophysical evidence

Figure 3.2.12. Map of magnetotelluric sites measured in the North German Basin between 1992-2004 (for more detail see Hengesbach 2007)

observed on the BABEL profile A that was the first survey to recognise the Moho and sub-Moho levels structurally in the area (BABEL Working Group 1993).

3.2.2.4 Lithospheric features from teleseismic investigations: Tomography and receiver functions Teleseismic methods are based on recording and analysing seismic signals from remote earthquakes. In the context of basin research they can be applied to investigate the gross structure of the crust and upper mantle and to define the major tectonic framework within which basin formation and development takes place. There are basically two teleseismic approaches which provide major information on the deep structure of the CEBS and the TESZ: seismic travel time tomography and receiver functions. The receiver function method is a powerful technique to gather deep seismic information on discontinuities in the crust and the upper mantle with three component seismic stations from teleseismic body waves (see Box 3.2.1-2 for detailed explanation). Since it has the lowest resolution of the different seismic imaging techniques we discuss here (see also Fig. 3.2.1), we use it to delineate the Moho and other large-scale structural features of the Palaeozoic base of the CEBS. Travel time tomography is based on analysing the arrival times of direct P- and S-waves with respect to theoretical travel times corresponding to an average “background” model of seismic velocities. The

output is a 3D field of seismic velocity “perturbations” indicating the modifications of the background model required to explain the observed travel times. As in seismic refraction, the interpretation of tomographic subsurface models is based on attributing changes of seismic velocity to changes in petrology, geological structure or temperature and pressure. In Central Europe, the TOR project (Teleseismic Tomography across the Tornquist Zone in Germany–Denmark–Sweden) recorded teleseismic signals between 1996 and 1997. It had a major impact on our knowledge of the lithospheric structure (e.g., TOR Working Group 1999; Gregersen et al. 2002). Deep lithospheric differences down to 300 km depth were detected and lateral boundaries defined (Fig. 3.2.11). At the northern rim of the Tornquist Zone and between the Elbe Line and the Caledonian Deformation Front, two marked lateral changes in seismic velocity are observed. At the Sorgenfrei-Tornquist Zone (STZ) the transition from fast P-wave velocities beneath the Baltic Shield to slower values below Denmark occurs within a short distance of about 50 km (Arlitt et al. 1999; Busche 2001; Shomali et al. 2002). It is associated with a dramatic increase in Moho depth (Gossler et al. 1999). The comparative figure of tomograms and receiver functions shows that the STZ can be traced down to 200 km depth and that it is also associated with a discontinuous structural change directly beneath the Moho (Fig. 3.2.11). At the Lower Elbe lineament (Elbe Line, EL) the receiver function image shows a strong lateral change in the lower

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Figure 3.2.13. 2D magnetotelluric inversion model of the Glückstadt Graben (modified from Hengesbach 2007). The eastern two-thirds of the profile correspond to the western part of the CRS stack shown in figure 3.2.7 and the central part of the potential field profiles in figure 3.2.22. Here, non-conductive sediments are found in the pre-Permian (yellow) and good conductors are denoted for the Carboniferous (greenblue). The geological interpretation (after Baldschuhn et al. 1996) is superimposed

crust and at Moho level which corresponds to the lateral zonation of the upper mantle velocities. Moho disruptions are observed in the area of the northern Variscan Deformation Front. Both receiver function images and seismic velocity tomograms show that the Trans-European Suture Zone comprises three distinct units, two of them underlying the CEBS. Towards the Glückstadt Graben area a regional thinning of the lithosphere from the Baltic Shield towards the Variscan Deformation Front is observed. The lithosphere/ asthenosphere boundary, indicated by low seismic velocities in the upper mantle, appears at 100 –125 km depth in the area extending between the Sorgenfrei-Tornquist Zone and the Elbe Line (Arlitt et al. 1999) and even shallower beneath the North German Basin towards the South. This thinning is not continuous but encompasses two pronounced steps: one beneath the Sorgenfrei-Tornquist Zone, the other around the Elbe Line (Gregersen et al. 2002; Plomerova et al. 2002; Shomali et al. 2006). The view that the Avalonia-Baltica Suture can be traced from the CDF down to the Elbe Line (Rabbel et al. 1995) is supported by the receiver function image as well as by the velocity increase in the uppermost mantle north of the Elbe Line. Across the Avalonia-Baltica Suture imaged in the NE German Basin along the BASIN9601 profile, the receiver function image confirms the results from steep- and wideangle measurements (see above; Krawczyk et al. 1999; Bleibinhaus et al. 1999) showing a flat lying Moho at ca. 30 km depth and the top crystalline basement as the strongest converting interfaces (Gossler et al. 1999). More diffuse conversions in middle crustal levels between 10–20 km depth are found predominantly between the Baltic Sea and

the Elbe Line. They are interpreted again to describe the transition from Proterozoic to Phanerozoic Europe.

3.2.3 Conductive layers and bodies from magnetotelluric observations Magnetotellurics (MT) is a passive electromagnetic technique to determine the conductivity structure of the earth at depths ranging from a few tens of metres to several hundreds of kilometres. MT rests upon the principle of induction, but unlike other electrical and electromagnetic methods it measures natural magnetic and telluric (electric) field variations, caused by solar radiation or atmospheric sources such as lightning discharge. Because of the so-called skin effect higher frequency ranges of the electromagnetic field give information on the shallow earth, whereas deeper information is provided by the lowfrequency range of the measured signal. The ratio of electric field to magnetic field is usually displayed as apparent resistivity or phase shift, and finally modelled or inverted. This yields information on the conductivity distribution in the subsurface. Good conductors or bodies may be imaged if coherent and especially connected fluids, melts or graphites along shear zones are present. For further reading, we recommend some very new textbooks with a vast number of references therein: Simpson and Bahr (2005); Burger et al. (2006); Spicak (2007). Many magnetotelluric sites and deployments along profile lines were measured in the last decade covering a large

Crustal structures and properties in the Central European Basin system from geophysical evidence

Figure 3.2.14. NE-SW trending magnetotelluric profile “Rügen Caledonides” across the NE German Basin and its northern margin (for location see Fig. 3.2.12). The good conductor between 5 to 12 km depth marks the trace of the shear-zone attributed to the Avalonia-Baltica Suture. It correlates with Cambro-Ordovician alum shales found in well G14 in the Baltic Sea (Piske and Neumann 1993). Along the profile, three wells and the MT station locations are marked

part of the North German Basin (Fig. 3.2.12; e.g., Hoffmann et al. 2001; Brasse and EMTESZ-Pomerania Team 2006). While many of those have contributed to the interpretation of regional structures (see Hengesbach 2007 and references therein), the major efforts first on a more regional scale covering the GG and then on a large scale with respect to the Avalonia-Baltica Suture are outlined here. The Triassic Glückstadt Graben (GG) is located above the deepest part of the North German Basin, where the Zechstein base reaches depths of more than 10 km (see also Fig. 3.2.7). The upper Rotliegend of the basin centre was predominantly composed of halites, which were squeezed into the abundant large salt walls during the Mesozoic. Only a relatively thin layer of clastics and volcanics remained (Brink et al. 1992). These rocks are abundant all over the North German Basin and do not show the seismic signature that is recorded below the GG. Therefore, the sedimentary strata of this unique seismic sequence were further investigated using MT measurements. 10 MT sites on an East-West orientated profile were set up across the total width of the GG (for location see profile Glück-stadt Graben in Fig. 3.2.12). The modelling results show that the central part of the graben is characterised by a good conductor at depths of approximately 10 km, which is missing on the adjacent graben flanks (Hengesbach 2007). The calculated depth is coincident with the depth of the Zechstein base (Fig. 3.2.13). The Kupferschiefer facies of the Zechstein base interval may have the potential to be a good electrical conductor; however, it is far too thin to be a layer that could account for the observed conductivity. Therefore, thicker sequences of organic rich layers have to be assumed. Based on nearby well results, black shales of the Early Carboniferous can be expected within

the graben proper and there only (cf. Brink et al. 1992 and references therein). Within the adjacent East Holstein area seismic reflections and refractions point to the presence of Early Carboniferous carbonates on the graben shoulders. This observation supports the assumption that the Glückstadt Graben has a Palaeozoic precursor and represents only a reactivated tectonic feature that had already been active during the ancient stages of the basin development. Focussing on the Caledonian Orogeny and its traces below the Southern Permian Basin, the available MT surveys allow the tracking of the Avalonia-Baltica Suture (ABS) which separates the older and stable Baltica craton from the younger and warmer Variscan-influenced Avalonia area further south (cf. Fig. 3.2.8). A good conductor that was calibrated close to the surface in the Baltic Sea as an alum shale horizon (e.g., drill-hole G14; Piske and Neumann 1993) is attributed to the shear-zone marking the collision of Avalonia and Baltica, buried today below the sedimentary NE German Basin (Fig. 3.2.14). This was also suggested by the interpretation of different types of seismic data in the NE German Basin (see above and also e.g., Krawczyk et al. 1999; Bayer et al. 2002), and was then additionally extended by MT surveys to other parts of the CEBS. In the example shown here, the sedimentary cover shows resistivities of max. 2 Ohm m, while the Permo-Carboniferous and older strata reach values of ca. 200 Ohm m. The ABS shear-zone can be traced from upper crustal levels down to depths of ca. 12 km as inferred from the good conductor of ca. 1 Ohm m (Fig. 3.2.14). This type of observation also occurs on other MT transects in the NE German Basin (e.g., Brasse and EMTESZ-Pomerania Team 2006). In contrast, zones of

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Box 3.2.2 Potential field modelling in 2D and 3D Equivalent masses and ambiguity The aim of gravity prospecting is to discover and study subsurface structures of deviating density by accurate measuring of the gravity field. When topography corrections have been applied a residual local variation of the Earth‘s gravity field remains, which can be attributed to the presence of masses. The study of the geometry of the bodies causing the observed anomaly is known mathematically as the Inverse Problem of Potential Theory. Problems arise because a given anomaly could be caused by many different bodies having the same density. For example, a uniform solid sphere produces exactly the same external gravity field as a concentric uniform spherical shell with the same total mass. It is not possible to distinguish these equivalent bodies by measuring only their external gravity field. Because of this ambiguity, the simplest approach to the inverse problem is to incorporate other constraints like e.g., seismic or other geophysical observations.

2D and 3D modelling We call a geological object two-dimensional if its length in one direction is infinitely long. Consequently all cross sections which are perpendicular to that direction have the same geometry and a constant density/susceptibility-distribution. However, the quantitative interpretation of real world geological bodies shows, that almost all bodies are three-dimensional. The mathematical formulation of 3D modelling algorithms is though, often difficult and results in complicated mathematical formulations. Without causing large errors we can call a body two-dimensional if its length (L) is four to five times its width (B). Modelling can also be scale dependent. As an illustration, from space the Andes are a 2D structure but as we move closer to the Earth‘s surface, they turn into a 3D structure.

Interoperable 3D modelling A 3D modelling program which can be used for forward modelling of potential fields is the IGMAS software (interactive gravity and magnetic application system; www.gravity.uni-kiel.de/igmas; Götze and Lahmeyer 1988; Schmidt and Götze 1998). The forward computation of the gravity and magnetic field of any geological model is based on the transformation of the volume integral to a surface integral. Therefore, the geological formations or volumes of constant density or susceptibility are approximated by polyhedrons with a triangulated hull. The 3D density model is set up by triangulation of polygons between vertical parallel planes. Because IGMAS provides also 3D GIS functionality profiles, faults and lineaments, bathymetry, gravity fields and other useful data and maps can be integrated in the model data base. The software package also provides interfaces to commercial 3D visualisation and GIS packages like GOCAD.

Crustal structures and properties in the Central European Basin system from geophysical evidence

3D Gravity backstripping methods 3D gravity analysis utilises a backstripping technique, whereby the calculated gravity effect of model layers, whose structure and properties are constrained by independent data, are successively removed from the observed gravity field (e.g., Starostenko and Manukyan 1987; Starostenko and Legostaeva 1998). One of the elementary bodies used for gravity calculations, is a vertical prism, where density varies with depth on linear or exponential functions. Thereby, the solution of the gravity problem is reduced to the determination of a double integral (Starostenko and Manukyan 1987). A prism is especially suitable to describe horizontally layered media. By combining a set of elementary prisms with inclined upper and lower surfaces it is possible to approximate complex 3D layering.

Curvature analysis One of the techniques to expedite the quantitative interpretation of gravity fields and to visualise lateral density changes is curvature analysis. Curvature attributes are surface-related and enhance linear aspects and/or properties of digitised surfaces, which may otherwise be difficult to observe (Roberts 2001). Toward this end, curvature calculations of gravity fields delineate areas of concave and convex curvature which expedites the detection of lineaments and faults.

Euler Deconvolution Euler deconvolution is a geophysical method intending to estimate the depth of gravity source points based on Euler’s homogeneity equation. Several authors have shown that Euler deconvolution leads to depth estimations of source points and that it can be applied to potential fields without additional constraints (Reid et al. 1990; Hoffmann 1999; Roy et al. 2000). The computation of regularised derivatives is performed in the spectral domain using the Fast Fourier Transformation and then multiplying by a regularisation filter for an optimal selected regularisation parameter (Pašteka and Richter 2002) to enhance the depth estimation.

good conductivity in the NW German Basin are attributed to Rhenohercynian alum shales (Early Carboniferous) or Westphalian coal seams (Late Carboniferous). These are found at 5 to 12 km depth with thicknesses varying between 1 to 4 km. Thus, the good conductors in the NW German Basin are interpreted in terms of rock type and metamorphic processes, while the conductors in the NE German Basin are interpreted tectonically.

3.2.4 Rock properties and density structure from potential field investigations Gravity and magnetic fields, also referred to as potential fields, are used to look beneath the Earth’s surface or more precisely beneath a certain reference level by sensing the

Figure 3.2.15. Left: Total magnetic anomaly of Northern Central Europe (after Wonik et al. 2001). Major lineaments separating Laurentia, Avalonia, Baltica and Variscan Europe are marked: CDF – Caledonian Deformation Front, VDF – Variscan Deformation Front, IS – Iapetus Suture. White line – Elbe Line, black line – gravity lineament at Tornquist Suture Zone. Right: The dip curvature map of the total magnetic anomaly in the direction of maximum dip enhances linear trends

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Figure 3.2.16. Map of the observed gravity field of the Central European Basin System (data base of Wybraniec et al. 1998). Major structural features are CG – Central Graben, GG – Glückstadt Graben, HG – Horn Graben, RT – Rheinsberg Trough, TTZ – Tornquist-Teysseire Zone, STZ – Sorgenfrei-Tornquist Zone, EFZ – Elbe Fault Zone, DSHFZ – Dowsing-South Hewett Fault Zone, CDF – Caledonian Deformation Front, VDF – Variscan Deformation Front, SNF – Sveconorwegian Front, MDH – Magdeburg-Dessau High, HL – Heligoland Low, PH – Pritzwalk High, RFH – Ringkøbing-Fyn High, SH – Sylt High

physical rock properties density and magnetisation respectively. Observations of the gravity field are made with the aim of determining the density and geometry of crustal geological bodies. In most cases, field observations of the gravity field itself are used to conduct interpretations at quantitative and qualitative levels. Lateral resolution of crustal density inhomogeneities is defined by the spacing of gravity stations. Potential-field surveys are relatively inexpensive and can quickly cover large areas of ground at the Earth‘s surface, on ships and aircrafts and most recently on satellite platforms. Gravitation is the force of attraction between two bodies, such as the Earth and a sensor and therefore a vector composed of the two components attractional and centrifugal force. The strength of this attraction depends on the mass of the two bodies and the distance between them. Gravity changes associated with many known and measurable factors such as movements of the vertical rotation axis, mass distributions after earthquakes and/or tidal forces. Higher gravity values are found over geological formations that are more dense, and lower gravity is found over rocks that are less dense, e.g., sedimentary basins. For modelling purposes in sedimentary basins down to depths of a few km (e.g., to the Zechstein horizon in the CEBS) the conversions of seismic velocities into rock densities are rather successful and provide good results (e.g., Gardner et al.

1974; Ludwig et al. 1970). More sophisticated approaches have to be applied in cases were the pressure, temperature and chemical rock composition of lithospheric domains prevent a simple conversion. In contrast to the gravity and magnetic fields, their gradients respond more sensitively to lateral density variations. Therefore, gravity gradients have frequently been used to detect and investigate fault systems and gravity lineaments (e.g., Blakely and Simpson 1986; Thurston and Brown 1994). Horizontal gradients were quantified by calculating gravity differences between different gravity observation points. Data and maps can be processed by forward and backward methods. Their differences are explained in box 3.2.2. Comprehensive articles and text books are available for potential field methods, further reading of which is strongly recommended: Vassiliou (1986); Arabelos and Tscherning (1987, 2001); Blakely (1996); Telford et al. (1976); Li and Götze (2001); Hackney and Featherstone (2003); Fowler (2004); Lowrie (2007). The potential fields of the Central European Basin System (CEBS) are composed of the signals of the more than 10 km thick sedimentary cover and of the heterogeneous structure of the deeper lithosphere below the basin (see

Crustal structures and properties in the Central European Basin system from geophysical evidence Figure 3.2.17. Separation of the residual gravity field into (a) its long wavelength component calculated by a low-pass filter of wavelengths below 200 km and (b) its short wavelength component calculated by subtraction of the long wavelength part (a) from the residual anomalies (after Yegorova et al. 2007). Abbreviations of tectonic units: Norw.-Dan. Bas. – NorwegianDanish Basin North-Germ. Bas. – NorthGerman Basin Polish Bas. – Polish Basin TTZ – Tornquist-Teysseire Zone STZ – Sorgenfrei-Tornquist Zone EEC – East-European Craton EFZ – Elbe Fault Zone DSHFZ – Dowsing-South Hewett Fault Zone CDF – Caledonian Deformation Front R-FH – Ringkøbing-Fyn High, VDF - Variscan Deformation Front SNF – Sveconorwegian Front CG – Central Graben GG – Glückstadt Graben HG – Horn Graben RT – Rheinsberg Trough

also under 3.2.1 above). Attempts to separate e.g., gravity signals from different depth levels in the lithosphere were made during the last decade, including several 3D regional gravity studies (Hermes 1986; Bachmann and Grosse 1989; Dohr et al. 1989; Brink et al. 1994; Zhou and Thybo 1997; Dirkzwager et al. 2000, Thybo et al. 2006; Yegorova et al. 2007). Most of these studies were primarily aimed at revealing the gravity signature of pre-Zechstein formations using gravity stripping (for method see Box 3.2.2). Magnetic field anomalies of the CEBS were compiled within the framework of the European Geotraverse project (EGT) and base on the reprocessing of magnetic field data from several European countries (Wonik et al. 2001). The dataset has been upward continued to 3000 m height and interpolated on a 5 km x 5 km grid (Fig. 3.2.15). Three deep crustal domains are identified: (i) mostly strong positive magnetic field anomalies N and NE of the Caledonian Deformation Front (CDF), (ii) intermediate to low amplitude negative and positive magnetic anomalies between

the CDF and Elbe Line (EL), and (iii) low small negative anomalies S and SW of the EL, interpreted as Avalonia crust. The dip curvature representation of the magnetic field emphasises lateral changes and the structural variability along the CDF (Fig. 3.2.15). The gravity field of the CEBS is characterised by a variable pattern (Fig. 3.2.16). The combination of Bouguer anomalies onshore and free-air anomalies offshore (Wybraniec et al. 1998; Banka et al. 2002; Bayer et al. 2002; Williamson et al. 2002; Yegorova et al. 2007) reveals two regional gravity lows over the Norwegian-Danish Basin and the North-German Basin, which are separated by the positive anomaly of the Ringkøbing-Fyn High. This first-order trend, however, is partially overprinted by anomalies of shorter wavelengths. These long and short wavelength components can be separated by the method of gravity-stripping which is suitable when the sedimentary cover is reasonably well

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Figure 3.2.18. Vertical profile 27 of the density model (left) and model statistics (right) showing the differences between modelled and measured free-air anomaly. Red curve - measured free-air gravity, black curve - modelled field, coloured lines - seismic reflectors and refractors serving as further geometrical constraints

Figure 3.2.19. 3D colour-coded density model and interpreted MONA LISA profiles 1 and 2 (black, after Abramovitz et al. 1998; Abramovitz and Thybo 2000), showing the morphology of the Caledonian Deformation Front at depth

Crustal structures and properties in the Central European Basin system from geophysical evidence

Figure 3.2.20. Left: Residual gravity anomaly at the base of the Zechstein horizon, Right: dip curvature map, indicating the transition from Baltica to Avalonia crust. CDF – Caledonian Deformation Front, EL – Elbe Lineament, GL – gravity lineament; green lines– graben bounding faults

known. For the CEBS, several main steps were undertaken (Yegorova et al. 2007), including (i) removal of the gravity effects of the sedimentary fill down to the base of Permian; (ii) subsequent removal of the gravity effect of the Moho topography; (iii) removal of large-scale crustal heterogeneities afterwards; and finally (iv) the separation of the resultant residual field into remaining long wavelengths and short wavelengths components (Fig. 3.2.17). The results suggest that the gravity effect of the Moho relief compensates to a large extent the effect of the sediments in the CEBS and in the North Sea. Removal of the effects of large-scale crustal inhomogeneities shows a clear expression of the Variscides at the southern part of the study area and the old crust of Baltica further NE (Fig. 3.2.17). The long wavelength anomalies are suggested to be caused mainly by density variations in the upper mantle, though gravity influence from the lower crust cannot be ruled out. Thus, the sub-basins of the CEBS originated on different lithospheric domains. The Polish Trough originated on the thick, strong and dense lithosphere of Baltica, the Norwegian-Danish Basin formed on a weakened Baltica low-density lithosphere formed during the Sveconorwegian Orogeny, and the major part of the North German Basin is characterised by a high-density lithosphere, which includes a high-velocity lower crust (relict of Baltica) overthrust by Avalonia (see above). The short wavelength pattern of the final residuals clearly displays several NW-trending gravity highs between the Tornquist Zone and the Elbe Line (see also Scheck et al. 2002a), e.g., south of the Ringkøbing-Fyn High the chain of positive anomalies of the Central North Sea, the Central

Graben, the Horn Graben, the Glückstadt Graben and the Rheinsberg Trough (Fig. 3.2.17). For 3D gravity modelling in the southern part of the North Sea, the gravity database consisting of three datasets was homogenised for forward modelling purposes and numerical field analysis e.g., spectral analysis, curvature and Euler deconvolution (see Box 3.2.2). The free-air gravity anomaly of the southern North Sea is characterised by WNW-ESE trending negative anomalies (up to -45x10-5 m/s²; Fig. 3.2.16). In addition to the stripping procedure described above, a 3D density model was compiled in order to separate gravity effects of the basin, the Variscan and Caledonian influenced sedimentary cover and the crystalline basement. The model is set up by 62, SW-NE directed vertical sections, thereby guaranteeing a high resolution even of superficial crustal structures (e.g., salt domes and walls), and by 8 horizontal layers, including the water body of the North Sea and the Zechstein. Model constraints came from bathymetry and reflection seismic data (TGS NOPEC, Norway; see also in 3.2.1 above and figure 3.2.2 for location). Vertical section 27 out of the 63 cross sections of the density model is shown in figure 3.2.18 which demonstrates the modelled crustal details and the mapped consistency between modelled and mea-sured gravity field. In the south and west of the model the crystalline crust is compiled as a 2-layered Avalonia crust (green units, Fig. 3.2.18) while the Baltica crust in the northeast is built by a 3-layered structure (pink units, Fig. 3.2.18). This density interface marks the trace of the Avalonia-Baltica Suture zone below the North German Basin, constrained by seis-

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of the regional gravity high while the gravity lows of the Central and Horn Grabens were removed (Fig. 3.2.20). The dip curvature of this field significantly enhances the characteristic gradients of which the most remarkable one is the NW-SE striking gradient west of Sylt (labelled GL; Fig. 3.2.20b). It parallels the CDF and is interpreted as the transition zone between the high-density Baltica and the low-density Avalonia crust (cf. Caledonian Suture Zone by Lyngsie et al. 2006). Small-scale features might indicate salt and diapiric structures. Further to the east, the Glückstadt Graben provides a study object that has to cope with the gravity effects resulting from large salt structures. These usually provide a density deficit in the observed field that can reach values of some 30 x 10-5 m/s2 and can therefore hide local gravity (Yegorova et al. 2007). To study and separate these effects, 3D gravity backstripping was performed within the area covered by the 3D structural model of the Glückstadt Graben and adjacent areas (Fig. 3.2.21; Maystrenko et al. 2005b, 2006; Yegorova et al. 2008).

Figure 3.2.21. Maps of the (a) gravity anomaly (Bouguer anomaly onshore, free-air anomaly offshore) and (b) total magnetic field, measured by an aeromagnetic survey at a height of 700 m above the Glückstadt Graben (after Yegorova et al. 2007a; Maystrenko et al. 2005b). Structural elements: CGG – Central Glückstadt Graben, EHT – Eastholstein Trough, HT – Hamburg Trough, WHT – Westholstein Trough. Elements of the gravity field: GL – Glückstadt Low, HAH – Hamburg High, HAL – Hamburg Low, HH – Heide High, HL – Helgoland Low, ML – Mölln Low, NKH – NorderstedtKiel High, RFH – Ringkøbing-Fyn High, SH – Sylt High

mic data (e.g., Thybo 1990; Rabbel et al. 1995; Thybo 2001). Its geometry was derived from seismic investigations along the MONA LISA profiles (see Abramovitz et al. 1998; Abramovitz and Thybo 2000; Domaschk 2000) and from findings of gravity processing studies. The overlay of seismic data and gravity model along these lines supports the geometry and areal extent of the Avalonia-Baltica Suture and its position and morphology southwest of the CDF (Fig. 3.2.19; see also in 3.2.1 above). In addition, the residual gravity anomaly at the base Zechstein horizon now shows a more continuous shape

3D gravity backstripping at the depth level of the base Zechstein shows that the basin gravity effect is very much influenced by the thickness of Zechstein salt so that short and long wavelength components are difficult to separate (Yegorova et al. 2007a). However, the gravity effect of Meso-Cenozoic sediments and Permian salt yields the most pronounced anomalies in the Westholstein Trough (-100 x 10-5 m/s2) and in the Eastholstein Trough (-85 x 10-5 m/s2). Over the axial part of the Glückstadt Graben, the gravity effect is in the range 70 – 80 x 10-5 m/s2. The total effect decreases gradually towards the basin flanks according to observed thinning of sediments in the same direction (Fig. 3.2.21). Since the residual field of the south-eastern part of the basin centre is characterised by a relative gravity low which corresponds to a broad Glückstadt Low in the observed field, this gravity low cannot be explained by the effect of Meso-Cenozoic sediments and Permian salt only. The crustal density anomalies are probably caused by older sediments and/or low-density crystalline crust within this part of the basin. The Heide Gravity High, as the most prominent positive anomaly of the Glückstadt Graben (Fig. 3.2.21), has a deeply rooted source as was shown by filtering the residual gravity field. Furthermore, this anomaly is also traceable in the short wavelength component. This supports former findings of gravity modelling within the Glückstadt Graben (Dohr et al. 1989) which indicate the presence of an intrusion within the crust beneath the Heide High. Derived from the 3D structural model of the GG (Maystrenko et al. 2006), two 2D density models in two cross-sections were analysed (Yegorova et al. 2007a; Fig. 3.2.22). The sedimentary cover of the Glückstadt

Crustal structures and properties in the Central European Basin system from geophysical evidence Figure 3.2.22. Calculated and observed gravity and magnetic fields along two 2D density modelled profiles across the Glückstadt Graben (for location see Fig. 3.2.21). Different density layers of the basin and lithospheric domains are shown by shadings and their densities by numbers in kg/m3. Magnetic domains and their susceptibilities are shown in SI units. The magnetotelluric profile in figure 3.2.13 covers the central part of profile 1

Graben is characterised by a low density (2.05 Mg/m3) within the upper part of the Cenozoic where sediments are relatively young, and therefore have a low degree of compaction. However, the Buntsandstein and the Muschelkalk successions reach a depth of more than 10 km in the basin centre and consequently the Early-Middle Triassic deposits are overlain by a thick sequence of younger sediments which cause a high degree of compaction in the Central Glückstadt Graben and the marginal troughs (the Westholstein and the Eastholstein Troughs). The density of salt structures varies from 2.18

to 2.35 Mg/m3 along the profiles (Fig. 3.2.22). The results of gravity modelling indicate that the Permian salt is not homogeneous due to differences in the degree of salt saturation in salt-rich bodies, and hence due to the contribution of Rotliegend salt associated with alteration with shale. There is a tendency that the highest densities of salt bodies (2.3 – 2.35 Mg/m3) occur in the huge salt walls within the central part and marginal troughs (Eastholstein, Westholstein and Hamburg), which have undergone the most complicated evolution with several pulses of salt movement during post-Permian times.

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Model structures are constrained by seismics and indicate a coherent reflectivity within the upper crust below the basin centre (e.g., Yegorova et al. 2007a; Maystrenko et al. 2005a,b). The rather schematic interpretation shows a lowdensity body along both profiles at depths of about 18 km (Fig. 3.2.22). The 2D density modelling indicates that the Glückstadt Graben is underlain by a high-density lower crust body with a density of 3.0 Mg/m3. The shape of the high-density body is in agreement with the results of 3D gravity backstripping which indicates a high-density lower crust beneath the entire Glückstadt Graben with local gravity maxima over the Westholstein and Eastholstein Troughs (Yegorova et al. 2007a). Most probably, the formation of

this high-density body in the lower crust is the result of the complex poly-phase tectonic history of the study area. The results of magnetic modelling across the GG indicate the presence of an area with high susceptibility (0.0850.145 SI; Fig. 3.2.22) mainly in the crystalline crust below the Central Glückstadt Graben. During modelling, a constant Curie-isotherm was considered at a depth of 22 km in accordance with previous magnetic studies within the North German Basin (Lindner et al. 2004; Scheibe et al. 2005). According to the results of magnetic modelling, the shape of the area with high magnetisation is not symmetrical. In addition, the results of 2D magnetic modelling

Figure 3.2.23. Three representative cross sections through the final 3D density model of the Bramsche anomaly (after Bilgili et al. 2007). The 3D model is built of seven geological units: Quaternary and Tertiary (yellow), Cretaceous and Late Jurassic (green), Early Jurassic (light cyan), Triassic and Zechstein (cyan), Cambrian/Silurian (gold red), Precambrian (dark red) and the intrusion (red). Numbers show densities in Mg/m3

Crustal structures and properties in the Central European Basin system from geophysical evidence

Figure 3.2.24. 2D cross-section from 3D gravity modelling in the NE German Basin (after Kuder 2002). An Avalonia accretionary wedge on top of Baltica crust (right half of profile) and Variscan influenced areas (left part of profile) are evident

indicate that the central part of the high-density body is overlain by an area with high susceptibility (0.08-0.145 SI). An unsolved problem of regional importance for both the evolution and structure of the Northwest German Basin is the existence or non-existence of the so-called Bramsche Massif (for location see Fig. 3.2.15). The high thermal maturity in the southern part of the Northwest German Basin (Petmecky et al. 1999), the CO2-risk and lack of reservoir rocks are regarded as the consequence of an intrusion that forms the Bramsche Massif. The rocks of the massif are believed to have intruded into the upper crust during the Late Cretaceous (Bachmann and Grosse 1989). The Bramsche gravity anomaly (Fig. 3.2.15) is known for decades. In recent years, Brink (2002b) has questioned the former interpretations of the Bramsche positive gravity anomaly. The strongest discrepancy is

that the level of hydrocarbon maturity measured in boreholes is better explained by an increased amount of burial with subsequent tectonic inversion. Additionally, a Vp/Vs ratio of 1.72 is not representative of a gabbroic intrusive body, but more likely of a dolomitic layer, here perhaps from Devonian/Carboniferous times. The density modelling was constrained by station-complete Bouguer gravity anomalies calculated from data provided by the GGA institute in Hannover. The dataset used for this study comprises about 4900 gravity stations and covers an area of 150×150 km. The gravity map of the region studied shows a slightly positive Bouguer anomaly of 10 x 10-5 m/s2 in the region: the Bramsche area is characterised by a positive Bouguer anomaly of up to 34 x 10-5 m/s2. This positive Bouguer anomaly also correlates with a magnetic high of 140 nT.

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The differences between the modelled and measured field reflect the missing mass related to the absence of an intrusive body. A good fit between measured and modelled gravity is only achieved in the models that include highdensity intrusive bodies in the Pre-Zechstein (Fig. 3.2.23). Incorporating an intrusive body at a depth of about 10 km yields a good correlation with the geometry of this body also constrained by the Euler source points (correlation 85%, standard deviation 4.86×10-5 m/s2). The final and best-fit model incorporates constraining information from seismic profiling and borehole density measurements. Here, we prefer a model that includes an intrusive body or, at least, a high-density body (Bilgili et al. 2007). Only such models cause a gravity field that fit well with measured data. In our final and preferred model, the distribution of Euler source points is used to constrain the shape and location of an intrusion. However, the gravity modelling does not allow conclusions with respect to the age of the modelled intrusion (Permian or Late Cretaceous). A Permian intrusion would be possible according to both gravity modelling and maturity modelling (see Littke et al. this volume; Voigt et al. this volume). Despite the good fit between measured and modelled gravity, further advances in the understanding of the crustal structure in the area of the Bramsche Massif require more detailed investigations, in particular from a seismic reflection profile. A similar anomaly feature is evident in the area of Pritzwalk in the NE German Basin (Fig. 3.2.16) where geophysical studies were conducted within the framework of the DEKORP-BASIN’96 survey (see above in 3.2.1). In these studies, the Pritzwalk anomaly was attributed to both a pluton intruded at 6-10 km depth and to an uplifted Moho. Hoffmann and Stiewe (1994), Bayer et al. (1999) and Brink (2002b) assumed that the anomaly is caused by lower crustal structures. On the basis of 3D gravity and magnetic modelling, Kuder (2002) showed that the Pritzwalk gravity high could be interpreted as the result of an uplifted Moho and/or uplifted lower-crustal structures and a high-density body (Fig. 3.2.24, profile km 200; for location see Fig. 3.2.8). On a more regional scale, the gravity modelling along the DEKORP-BASIN 9601 profile proposed a high-density lower crust below the centre of the NE German Basin and an increased thickness of low-density upper crust below the margins to fit observed and modelled data (Kuder 2002; Fig. 3.2.24). Geodynamically, such a supply of high-density material could be due to the transfer of mantle material into the crust by magmatic underplating. Indications corroborating this hypothesis are also provided by seismic data (Rabbel et al. 1995; Krawczyk et al. 1999; see above in 3.2.1), studies on magmatic activity (Benek et al. 1996) and subsidence analysis (van Wees et al. 2000; Scheck et al. 1999).

The continuation of Baltica below the North German mainland has also been modelled by gravity data, which suggest the presence of an Avalonian accretionary wedge on top of Baltica crust that thins out towards the NE German Basin depocentre (Kuder 2002; Fig. 3.2.24). An alternative interpretation is that there may be a low-density body below the northern margin of the Southern Permian Basin. This, however, is a less favourable model since reprocessing of seismic data in the region has indicated the presence of an accretionary wedge (see above; McCann and Krawczyk 2001; Lassen et al. 2001; Krawczyk et al. 2002, 2008).

3.2.5 Summary The CEBS rests in the N on Caledonian- and in the S on Variscan-influenced basement: thus its underlay comprises a number of terranes derived from the three continental landmasses Laurentia, Baltica and Gondwana (with Avalonia as the northernmost terrane) which collided during Ordovician-Silurian times to form the basement of northern Europe. The geophysical portrayal of the nature of the crust and upper mantle by deep seismic and potential field studies at the present day includes all of the postCaledonian changes which have occurred to modify its characteristics. Such changes include salt tectonics in a regional and local context (see Kukla et al. this volume), as well as events related to the Variscan Orogeny, Mesozoic basin evolution, Alpine deformation transmitted into the foreland and other Cenozoic extensional events (see Kley et al. this volume; Sirocko et al. this volume).

·

Main results from the recent geophysical work are:

· · · ·

In the North Sea, salt diapirs formed predominately in close proximity to the major grabens, but did not originate close to the West-Schleswig Block. This depends possibly on the initial thickness of salt, the depth and thermal differences between both locations and zones of crustal weakness. Touchdown-structures are common, where younger sediments directly overlie the pre-salt sediments in the North Sea while salt is missing. This may be related to either basement fault activity or just depletion. The Glückstadt Graben was already an important depression in the Early Carboniferous. Late CretaceousEarly Tertiary compression caused increased salt activity. A flat lying Moho does not support the classical extension model, or it may indicate a new phase of petrophysical equilibrium. The Post-Permian supra-salt successions in the Baltic Sea show a thin-skinned tectonic response to subsequent tectonic events. New salt pillows are detected in the Baltic Sea (Rerik area). In contrast to major graben and rifting structures in the

Crustal structures and properties in the Central European Basin system from geophysical evidence

western part of the Southern Permian Basin, a growing influence of the cold and stable Baltica craton is observed towards the eastern basin parts. Despite these major efforts, some debate still remains, addressing for example questions like: To what extend does the reactivation of faults contribute to the neo-/tectonic activity, and how is it related to crustal heterogeneities? How important were strike-slip movements? What was the subduction polarity during the Avalonia-Baltica collision, or was there even a flip in subduction polarity? Do the observed lineaments exist below the lithosphere-asthenosphere boundary?

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In summary, the aforementioned articles, comprehensive reviews and selected results, together with the

geological maps of Ziegler (1990), the newly compiled “Geology of Central Europe” (McCann 2008) and the upcoming Petroleum Geological Atlas of the Southern Permian Basin Area, palaeogeographical reconstructions via the PALEOMAP project (Scotese 1999), potential field data by the British Geological Survey (http://www.bgs. ac.uk/geophysics/home.html; Pedley et al. 1994), Banka et al. (2002) and Wybraniec et al. (1998), as well as the Geothermal Atlas of Europe (Hurtig et al. 1992) provide a very thorough geophysical database for the Central European Basin System. In addition to the initial mapping character of some large-scale surveys, the combination of the different geophysical methods together with the geological knowledge exploits only the full potential of each attempt. Finally, these results can be handed over to structural model building, geodynamic modelling or site planning.

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Chapter 3.3

Strain and Stress J. Kley · H.-J. Franzke · F. Jähne · C. Krawczyk · T. Lohr · K. Reicherter · M. Scheck-Wenderoth J. Sippel · D. Tanner · H. van Gent – the SPP Structural Geology Group*

3.3.1 Introduction This chapter deals with deformation structures in the Central European Basin system (CEBS) and how they evolved. The analysis of structures in any environment must proceed in three steps: beginning with the geometry of structural elements, followed by kinematics and finally dynamics (Box 3.3.1). The first two steps are necessary to determine the deformation, or strain, the third to determine the causal stress. Strain of the lithosphere at different scales is always involved in the formation of sedimentary basins. Even for those basin types where faulting does not play a major role – e.g., flexural foreland basins at some distance from the thrust front – the differential subsidence which created the basin in the first place is a first-order component of strain, expressed as regional gentle warping of the basin floor. In most basins, strain is also accommodated through displacements on major fault systems with associated block rotation and folding (Fig. 3.3.1). Smaller faults and folds often occur down to the outcrop scale at least. It has been argued that fault systems often exhibit fractal, scale-invariant geometries. How-

ever, there is a lower limit to the scale of deformation: penetrative strain, such as that causing slaty cleavage by (sub-)microscopic alignment and folding of clay minerals, is typically absent from sedimentary basins, except for very deep ones. The different methods used to analyse various components of strain will be outlined below. Stress (the sum of forces acting on a rock volume) is usually considered the cause of strain. Yet, in many cases strain and stress are not related in a simple manner (Fig. 3.3.1). The history of stress states over time – the dynamic evolution of a basin – is more difficult to unravel than that of strain states. Past states of stress leave no trace except for permanent strain increments they may have caused. In contrast, all geological deformation increments will add to the finite strain ultimately observed, except for small elastic strains that are recoverable. Hence, while presentday stresses can be measured or at least estimated using various methods, “palaeostress” analysis also starts with an analysis of the strain accommodated by systems of faults (see section 3.3.4 below). The information on the vertical strain component associated with subsidence and its evolution over time is stored in the sediment fill of a basin, particularly in thickness variations. Figure 3.3.1. Conceptual sketch of strain and stress components in an active extensional sedimentary basin. Strain (kinematic) components are labelled blue, stress (dynamic) components are labelled in red italics. Stress fields are indicated by the orientation of the principal normal stresses σ1, σ2, σ3 and by the attitude of the σ1-σ3 plane. A salt layer is shown to partly decouple the stress fields above and beneath it

*with contributions by R. Müller · J.L. Urai · T. Voigt

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Box 3.3.1 Geometry, kinematics, dynamics Basin modelling includes reproducing the structure of a basin or a part of it at a certain scale. The structural geometry can subsequently be analysed with respect to its kinematics (the way the geometry evolved over time) and dynamics (the changing forces or stress states that were the cause of the kinematic history).

Geometry The primary dataset constraining the geometry of stratigraphic contacts, faults and other fractures, is typically based on a combination of subsurface information from e.g., seismic interpretation and boreholes, and projected map and outcrop data. The amount, type and quality of data will determine whether a 3D or 2D model can actually be built. It is inevitable that these facts must also be augmented with some assumptions, since no dataset can be truly complete. The accuracy of the model and therefore its use is directly dependent on the accuracy of the original data and the methods applied. Seismic resolution depends on depth and acquisition; in the vertical it is on the meter scale (at the surface) to ca. 50 m (at 5000 m depth) for modern acquisitions. In the horizontal direction, the first Fresnel zone sets the resolution at a quarter of the wavelength of the seismic waves (Yilmaz and Doherty 2001). If the velocity model building is incorrect depth conversion will contribute errors. Borehole data are accurate to less than a metre, but they are one-dimensional. Outcrop data are two/three dimensional and as accurate as a borehole, but only at the surface. Projection of outcrop data to depth, even if confined within dip-domain boundaries, is possible, but the positional error increases with depth by the tangent of the angular error at the surface. Therefore, this method is only applicable to show the structure in the first ca. 1 km of the crust. Ideally, as many different types of data as possible should be combined, e.g., seismics, especially after depth conversion. Results should be verified with boreholes and combined with outcrop data in the first kilometre of the model.

Kinematics In principle, a full kinematic model would include the particle paths over time for all points in the volume under study. In practice, there are a number of ways to approximately define the kinematics of structures. Kinematic data include, e.g., in order of decreasing accuracy, slickensides in outcrop, the “a” direction of folds as determined by folded lineations (Ramsay 1969), the vergence of minor folds and the plunges of their axes, the topography of fault surfaces in outcrop and seismics (the direction of less roughness is the most probable movement direction), and large-scale fault curvature. While these data constrain the directions of particle motion, geometric modelling and structural restoration are required to determine the lengths of particle paths. Two-dimensional models must always be made in the (kinematic) transport direction of the major structures. Furthermore, plane strain (there is no movement of material in or out of the section plane) must be assumed. Three-dimensional models can avoid this assumption, but then require methods of restoration that can include divergent and convergent flow paths. In order to estimate the velocities of particle motion, the timing of deformation must be closely constrained. In sedimentary basins, the timing of fault motion or folding can usually only be determined indirectly. Unconformities of strata above the structure can give an indication of the time span of movement of the latter. Of course, the time indication is only as accurate as the stratigraphic resolution.

Dynamics Dynamic analysis is the determination of the forces or stresses (force per area) that have induced the particular kinematics of a rock volume. Present-day stress states can be determined using a number of methods (Box 3.3.4). Past stress states that have only led to relatively minor strain are accessible to palaeostress analysis (Box 3.3.3). Actual simulation of fold or fault dynamics in complex situations or involving large strains is only possible with analogue or numerical models. These models go beyond the purely geometrical representation of strain typical of kinematic models as described above. On the other hand, both analogue and numerical models have resolution problems that still preclude an adequate representation of the actual mechanical variations present in sedimentary basins. Analogue models also suffer from scaling problems, while numerical models often employ simplified rheologies (e.g., elastic behaviour), either for technical reasons or because no flow laws exist to describe the combined effects of different processes governing the deformation of rocks in sedimentary basins. At any rate, both types of models need true geological parameters as input to give meaningful results.

Regional thickness patterns provide a first-order proxy for vertical movements. Large thicknesses indicate local subsidence, whereas reduced or missing thickness indicates local relative uplift or stability. The load of the sediments imposes a component of vertical stress in itself that can amplify but also mask the tectonic signal. As sediment thickness is additionally influenced by exogenic factors like climate or depositional environment, it is useful to

separate the load induced component of subsidence from the total subsidence and to evaluate the remaining tectonic signal. This is commonly done using the method of “backstripping”. With this approach the total load of the sedimentary layers is calculated. The layers are then successively removed from the top downwards (backstripped), the palaeotopography is adjusted to palaeo water depth and the compaction of the underlying layers as well as the

Strain and Stress

Figure 3.3.2. Cross-section balancing methods. The lines whose lengths are kept constant during deformation are shown for each method, with one line highlighted in red. Further explanations in the text

respective isostatic equilibrium is recalculated according to the new load condition (Box 3.3.2). The distribution of the determined tectonic subsidence gives a more concise image of the vertical strain in response to tectonic stresses. Depending on the amount of data that can be used or handled, a backstripping analysis will in general reveal a regional picture and smooth over all but the largest faults. This method provides no information on horizontal strain components, i.e., extension or shortening. The nature and magnitude of the strains accommodated through movements on faults and folding can be estimated by structural restoration or balancing techniques. These involve unfolding of layers and restoring fault offsets to arrive at an undeformed configuration, in the simplest case flat-lying strata. The particle paths from the undeformed to the present state – if possible determined through various increments of the total deformation – define the kinematics of the deformation. Most balancing methods are based on the assumption that rock volume is conserved during deformation. Unlike the backstripping technique explained above, compaction is normally not taken into account. In a strict sense, classical balancing will therefore not work for synsedimentary deformation. However, if the volume change during deformation can be sufficiently well constrained, it is possible to combine backstripping and kinematic balancing techniques (see below and Box 3.3.2). Even when balancing is used for rocks in an advanced diagenetic state, processes such as pressure solution and precipitation may cause volume loss or gain during deformation. Nevertheless, the constant volume assumption often seems to be a reasonable approximation for deforming strata in sedimentary ba-

sins. Balancing was initially developed for two-dimensional problems, i.e., cross-sections (Dahlstrom 1969; Laubscher 1965). For cross-section balancing to be valid, not only the volume but also cross-section area must be conserved (plane-strain, see Box 3.3.1). This means that cross-section lines must be chosen with care, and that sections across strike-slip faults cannot be balanced at all. If the plane-strain criterion is met (again, approximately), and bed thicknesses do not change as a result of deformation, bed-length balancing can be applied: the lengths of stratigraphic horizons as measured between reference or “pin” lines at both ends of the cross-section must be the same before and after deformation (Fig. 3.3.2a). With classical balancing techniques, this is usually achieved through several cycles of trial-and-error. The difference in length between the deformed and restored cross-sections gives the magnitude of extension or contraction, Δl. For certain classes of fault-related folds, geometric models have been developed which link properties of the fold, in particular limb dips, to those of the associated thrust fault (Suppe 1983; Suppe and Medwedeff 1990). These models permit the construction of an arbitrary number of intermediate stages between undeformed and fully deformed, each stage being perfectly balanced. They are especially suited to forward modelling, i.e., starting with an undeformed state to arrive at the final configuration. Other cross-section balancing techniques were developed for cases where bed lengths do change (e.g., “inclined shear”, especially for extensional systems (Dula 1991); “trishear” at fault terminations (Erslev 1991) or to make for easier computation (e.g., “fault-parallel flow”). Many of these techniques also balance line-lengths, though not

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Figure 3.3.3. 2.5D and 3D structural restoration. (a) modified from Bitterli (1990). (b) from Wilkerson et al. (1991)

those of stratigraphic contacts, in order to maintain area balance (Fig. 3.3.2b, c). Attempts were made early on to extend structural restoration to three dimensions. Deformation magnitudes from serial balanced cross-sections can be used to restore map-view mosaics of fault-bounded blocks to their original positions, with the blocks being either rigid or deformable (Fig. 3.3.3a) (Bitterli 1990; Kley 1999; Laubscher 1965; Rouby et al. 1993). This technique yields information on strain gradients in map view and rotations about vertical axes, data that can otherwise only be obtained through palaeomagnetic studies. Closely spaced, strictly parallel sections permit modelling of pseudo-3D deformation with strain gradients along strike but no deviation from plane strain (Fig. 3.3.3b; Wilkerson et al. 1991). Other authors have developed tools to flatten out the surface of a single key layer (e.g., Gratier et al. 1991). Fully 3D restoration algorithms are already available in some computer programs, but are typically restricted to single structures and relatively simple structural geometries. Due to its particular mechanical properties, salt presents a major problem for structural restoration. Salt basically behaves like a viscous fluid over geological time spans and can change shape much more radically than other sedimentary strata. This makes estimates of its initial thickness difficult. The plane strain assumption will not hold for thick salt layers, as is evident from the radial movement of

salt into pillows and diapirs. In addition, when diapirs approach or pierce the surface, the volume of salt removed by dissolution is harder to quantify than the eroded volume of other strata. The usual way around these problems is to balance the layers above the salt and adjust the salt volume (or area) to the remaining space. Strata deposited during the rise of salt structures often give excellent clues to their progressive development (Trusheim 1957). Since thick salt decouples structures above and below it, problems often remain when attempting to restore the elevation and structures of the basement beneath the salt. In the best case, data coverage permits the construction of 3D models that can be sequentially restored. However, restoration of well-positioned 2D sections may yield very helpful insights into the deformation mechanisms that have shaped a certain area. Possible restoration workflows are shown in box 3.3.2. These workflows ideally combine kinematic reconstruction with isostasy and decompaction and require knowledge of the geometry, lithology and related distribution of physical properties such as density, porosity, compaction behaviour, as well as information on palaeo water depth for different time steps. Estimates of volume loss due to dissolution and/or erosion of layers may further constrain the restoration. In real case studies at least some parts of this required information may not be available and must be replaced by assumptions. The presence of salt in a basin poses additional difficulties for

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Box 3.3.2 Restoration techniques Backstripping

Sequential restoration of basin deformation

is a method used in subsidence analysis to separate the loadinduced, isostatic component from the total subsidence and to determine the tectonic subsidence. The individual layers (A-D) are removed sequentially. For each backstripping step n, the uppermost layer is removed and the new isostatic equilibrium, as well as the new compaction of the remaining layers, is calculated according to the change in sediment load and palaeo water depth (PWD). The difference between total and load-induced subsidence represents the amount of subsidence due to tectonic factors (horizontal stress, thermal cooling). In the simplest case, local (Airy) isostasy is assumed:

helps to unravel the mechanisms behind observed subsidence and uplift. Though different tools and methods are available, the restoration workflow may consider one or several physical processes, depending on available data quantity and quality. Subsidence restoration should consider the basic principles of backstripping. Fault restoration and unfolding may be performed using different kinematic approaches.

ρcrust*mcrust=ρmantle*mmantle (ρ=density, m=thickness) where ρcrust*mcrust may consist of n layers with individual thickness and densities. Removal of a layer n will change the load ρcrust*mcrust. This change of the component ρcrust*mcrust results in a change of (palaeo) topography. (If ρmantle=const, then mmantle has to change accordingly).

3D Backstripping and salt restoration can be used to derive the history of salt movements over time. Basic assumptions are that the salt behaves like an incompressible viscous fluid and moves in response to differential loading (Scheck et al. 2003b).

a. The pressure at the base of the salt, P, is the sum of the pressure acting on the salt upper surface Ptop and the load of the salt column gρshs.

b. Under hydrostatic conditions, the pressure difference ΔP is zero within the salt between two points 1 and 2, at the same depth zn:

P=Ptop + gρshs

ΔP[1,2]= 0 = = P1-P2 + gρs (h1-h2)

c. After removal of a layer, ΔP [1,2]≠0! =>the salt has to be redistributed until the condition is true again. =>Introduction of a formal flux of salt “j” that takes place to equilibrate the system with: j=Δ(Ptop+gρshs+gρsZbase)

d. Once hydrostatic conditions at the base salt are re-established, isostatic restoration in response to total change in load (ρsed*msed) needs to be considered. ρsed*msed+ρcrust*mcrust= = ρmantle*mmantle

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restoration as neither plane strain nor uniform mechanical properties can be assumed. Though different attempts to restore salt-tectonic deformation are still under debate, one possible strategy is to consider differential loading of the salt as the main effective mechanism (Box 3.3.2).

3.3.2 Structural framework of the CEBS Much of the first-order structural configuration of the CEBS can be read from bulk sedimentary thickness maps (Lamarche and Scheck-Wenderoth 2005; Scheck-Wenderoth and Lamarche 2005; Ziegler 1990; Maystrenko et al. this volume; Scheck-Wenderoth et al. this volume; Littke et al. this volume) which reflect long-term integrated basin subsidence (Fig. 3.3.4). Of the three main sub-basins defined within the CEBS, we will here focus on the North German Basin (NGB), but will also briefly refer to the NorwegianDanish and Polish basins and to certain areas outside the

CEBS. This is necessary because some structural systems are not confined to the present-day basin outlines, indicating that deformation occurred outside the basins or that erosion has obliterated parts of the original basins from isopach maps. The following overview of the structures within the CEBS or related to it, draws on the compilations by Baldschuhn et al. 2001, Lokhorst et al. 1998, Mazur et al. 2005, and other sources. The degree of detail and interpretation varies considerably between these different sources. A full map-view interpretation of varying fault activity and nucleation for different time slices is only available for part of Northern Germany (Baldschuhn et al. 2001). Of all the structures shown in Fig. 3.3.4, only some major ones can therefore be confidently assigned to certain time intervals. Nevertheless, we will make an attempt to describe the evolution of the general structural patterns over time. The present-day configuration of the CEBS (Figs. 3.3.4, 3.3.5) is dominated by NW-trending structures, parallel

Figure 3.3.4. Structural framework of the CEBS. Compiled after Baldschuhn et al. (2001), Mazur et al. (2005) and other sources. Only major faults are shown. Salt structures are ommitted for clarity (See figure 5.3.1 for a map of salt structures). Abbreviations are: Fault zones/ structures (black): STZ Sorgenfrey-Tornquist Zone, TTZ Teisseyre-Tornquist Zone; CG, HG, GG Central, Horn and Glückstadt grabens, Rh Rheinsberg Trough; RB Rheder Moor-Blenhorst, AT Allertal, Gi Gifhorn, Ga Gardelegen, Wi Wittenberg, LE Lausitz Escarpment, La Lausitz Thrust, Od Odra, Os Osning, St Steinhuder Meer, Eg Egge, FL Franconian Line, LN Landshut-Neuötting, Do Donau, Pf Pfahl. Subbasins (blue): SoPB Sole Pit, BFB Broad Fourteens, WNB West Netherlands, CNB Central Netherlands, LSB Lower Saxony, NEGB Northeast German. Basement highs (brown): RFH Ringkøbing-Fyn High; Ha Harz, Fle Flechtingen High, B. “Block”, HCM Holy Cross Mountains

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Figure 3.3.5. Structures of the North German Basin, after Baldschuhn et al. (2001) and other sources. N-S grabens and Late Cretaceous inversion structures are emphasised. Abbreviations as in figure 3.3.4. Additional abbreviations: RG – Roer Graben, LB – Leer-Bremen fault zone, We – Weser trough, TW – Thüringer Wald. Locations of figures 3.3.6, 3.3.13, 3.3.15, 3.3.16, 3.3.17 and 3.3.18 are indicated. Also shown are locations of case study areas. Western Allertal area comprises figures 3.3.9, 3.3.10 and 3.3.19

to the main axes of maximum total thickness. A second, less developed class of structures trends approximately N, varying between NNW and NNE. Structures of intermediate orientation between the two major trends or curving from one trend into the other also occur, while those trending NE are relatively rare. In many cases the N-trending faults link right-stepping segments of the longer NW-trending fault zones, making the overall pattern markedly asymmetric. The regional structural trends can be seen in faults, folds and salt diapirs or salt walls, and they can often also be observed at scales much smaller than the entire basin. A pattern of NW- and NNE- trending faults existed as early as Late Carboniferous time (Benek et al. 1996; Lokhorst et al. 1998). These faults formed a system of strike-slip and normal faults (e.g., Gast and Gundlach 2006; Mogensen 1995). The distribution of Early Rotliegend volcanics and 1

sedimentary basins shows that the fault system extended beyond the CEBS, and that large-offset faulting occurred outside it (e.g., in the Oslo graben and Saar-Nahe basin). Several of the early major fault zones were active later on, sometimes throughout the basin history. Examples include the NW-trending Tornquist Line (comprising the Sorgenfrey-Tornquist Zone (STZ) and Tornquist-Teisseyre Zone (TTZ)) and the NNE-trending Rheinsberg lineament1. The latter element flanks a swell, along which the Polish Basin bifurcates into the Norwegian-Danish and North German Basins around the eastern end of the RingkøbingFyn high (Scheck-Wenderoth and Lamarche 2005). The pattern of coexisting NW- and N-trending faults can still be observed on a smaller scale in late Rotliegend time. Upper Rotliegend strata are often quite pervasively dissected by synsedimentary normal and strike-slip faults at

Many of the fault zones discussed here are commonly termed “lineaments”. We use the more neutral term “fault zone” here, or “fault array” where faulting affects a wider swath. The names we use for individual fault zones are sometimes shortened, as in “Gifhorn array” instead of “Braunschweig-Gifhorn Lineament”. However, we have always retained part of the original name to avoid confusion. Abbreviations are used in figures 3.3.4 and 3.3.5. In the text, these abbreviations are cited in brackets once where the respective name is first mentioned.

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km-spacing. However the stratigraphic throws are relatively low, typically ranging from several tens to a few hundred meter. Faulting thus appears as a secondary phenomenon during this stage of the basin evolution, while subsidence was mostly due to lithospheric cooling after the Lower Rotliegend event of stretching and magmatism (van Wees et al. 2000). After Rotliegend time, the NGB can be subdivided into two large E-W-trending belts. The northern belt, adjacent to the Ringkøbing-Fyn High and comprising the “Pompeckj block” and Northeast German Basin (NEGB), is dominated by salt structures. It has comparatively few basement faults, which trend mostly around N. The southern belt has a much denser system of mostly NW-trending faults and comparatively few salt structures. The N- to NNE-trending normal faults of the northern belt formed or were reactivated in Triassic time, when a series of major grabens evolved. These grabens appear as an eastward extension of the North Sea rift. Going east from the North Sea Central Graben (CG) they include the Horn (HG) and Glückstadt (GG) grabens and the Weser (We) and Rheinsberg (Rh) troughs or depocenters. Only the Central Graben and, in a much more attenuated way, the Horn Graben dissect the Ringkøbing-Fyn High, whereas the remaining extensional systems were essentially confined to the area south of it. This suggests they must have been linked in the north by a NW-trending, dextrally-transtensional transfer fault. No continuous fault zone has been mapped there, but a conspicuous swath of left-stepping en-echelon faults runs along the southern margin of the Ringkøbing-Fyn High (Fig. 3.3.4). Other NW-trending fault zones within the NGB, e.g., the Allertal fault zone, also connect to N-trending grabens (Röhling 1991). In the Norwegian-Danish Basin and the Polish Basin, NWtrending depocentres prevailed in Triassic time. A major NNE-trending zone west and south of Bornholm connects the structures of the Danish and Polish sub-basins. The NW-trending faults of the southern belt originated as normal or transtensional faults from Late Triassic time onward (Baldschuhn et al. 2001), marking a southward shift in the focus of deformation. Large normal faults of Late Jurassic through Early Cretaceous age formed mostly in the western part of the southern belt, accentuating the Lower Saxony Basin (LSB) and the NW-trending grabens of the Netherlands and southern North Sea (Fig. 3.3.5). Judging from thickness variations and structures, the LSB was originally a large, somewhat asymmetric and internally faulted graben, bounded by the Osning (Os) master fault zone in the SW and by the Rheder Moor-Blenhorst (RB) and Allertal (AT) fault zones in the NE. Although the Zechstein salt decouples structures in the sedimentary cover from basement faults, all major suprasalt fault zones of the LSB are interpreted to be underlain by base-

ment faults (Kockel 2002). The Osning fault connects in the east with the Egge (Eg) system running toward the S, up to a considerable distance from the LSB. Normal faults and grabens splay off from the Egge fault system toward the SE and eventually appear to link up with the basement faults along the SW border of the Bohemian Massif. The northern border of the LSB is structurally complex as well. The Rheder Moor-Blenhorst fault zone and other normal faults within the LSB connect with the NNW to NW trending Steinhuder Meer (St) fault zone. East of the Steinhuder Meer fault zone, the more northerly located Allertal zone takes over as the northern border fault of the LSB. In the east, the LSB as defined by the presence of thick Late Jurassic-Early Cretaceous strata abuts on the transverse NNE-trending Gifhorn (Gi) fault array. This regional system already separated the Weser trough depocentre from the Eichsfeld-Altmark swell in Early Triassic time. Further east, the NEGB and westernmost Polish Basin structurally resemble the Pompeckj Block. They are nearly undeformed except for a few NNE-trending Triassic faults and salt structures. East of the Gifhorn system the present-day southern margin of the NGB and Polish Basin trends approximately east and is defined by an enechelon array of left-stepping, NW-plunging basement uplifts (the Flechtingen (Fle), Lausitz and East Sudetic uplifts and Holy Cross Mountains (HCM), Fig. 3.3.4). These uplifts do not represent the original basin margin but were mostly created later by tectonic inversion. Late Cretaceous to Palaeocene contraction, often reactivating the earlier normal faults, modified the NW-SE trending extensional structures right across the CEBS, but also created new thrust faults. Overall, contractional deformation affected a belt of relatively constant width, paralleling the Tornquist Line between the southern North Sea and the present-day Alpine-Carpathian front. The magnitude of deformation does not seem to vary much as one passes from the LSB eastward into the area of basement uplifts, although the structural style changes significantly. In the LSB, essentially all major Jurassic-Early Cretaceous extension structures are interpreted to have suffered some inversion (Baldschuhn et al. 2001), but reverse reactivation was most prominent along the former basin-bounding faults (Mazur et al. 2005). A large, flattopped basement uplift runs along the southern basin margin in the hanging-wall of the Osning fault (Figs. 3.3.5, 3.3.6). This fault system inverted the deepest part of the LSB (Senglaub et al. 2005; Adriasola-Muños et al. 2007) and partly carried it onto the southern basin margin. Being still largely covered by Triassic and Jurassic sediments, the associated basement uplift appears more subdued than its exposed eastern counterparts (see above), but its size and vertical offset are of the same order. As with Triassic to Early Cretaceous extensional features, the southern front

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of contraction steps to the right along the Egge system and continues along the SW border of the Bohemian Massif. A mirror-image, but shorter step to the left in westernmost Germany gives the Rhenish Massif the shape of an indenter into the southern border of the LSB (Fig. 3.3.5). The former basin-bounding Rheder Moor-Blenhorst and Allertal fault zones were reactivated as north-directed thrusts. However, inversion also affected fault zones on the Pompeckj Block (most prominently the Leer-Bremen (LB) fault zone; Fig. 3.3.5). Judging from its geometry, the Steinhuder Meer fault zone and its southeastern continuation formed a dextrally transpressive transfer zone between the Rheder Moor-Blenhorst system and the Harz basement thrust with its associated faults. East of the Gifhorn system the inverted LSB is replaced by a wide belt of thrust-related basement uplifts that extends southward to Bavaria and eastward to Poland. In contrast to the LSB which is thrust over its southern border, the basement blocks overthrust the NEGB and western Polish Basin and now define their southern margins (Figs. 3.3.5, 3.3.6 b) (Kossow and Krawczyk 2002; Otto 2003). The basement thrusts feed displacement into the Zechstein salt detachment and are thus linked to thin-skinned structures in their foreland. Salt structures and remnants of Jurassic strata on top of and between the basement uplifts suggest that the

area of basement thrusting was originally not much different from the remainder of the Northeast German Basin. Reverse faulting coincided more closely with the extensional basin outlines along the Tornquist Line from Norway to Poland and provides many classical examples of inversion tectonics (Deeks and Thomas 1995; Krzywiec 2006b; Mazur et al. 2005; Mogensen and Jensen 1994). Only a few and relatively minor structures of the CEBS formed after the contractional deformation in latest Cretaceous or early Tertiary times. Existing faults were in some cases reactivated as a result of salt movements or mild extension (Kockel 2003). The Central European Rift System affected the area only marginally. However strong Cenozoic subsidence in the southern North Sea, apparently not associated with major faulting (Hall and White 1994), induced a gentle regional northward tilt over much of the NEGB.

3.3.3 Structural analysis and quantification of strain In this section, we first present a conceptual outline of deformation processes associated with the inversion of the

Figure 3.3.6. Two regional cross-sections, located west (a) and east (b) of the Gifhorn fault array. No vertical exaggeration. Cross-section locations in figure 3.3.5. Notice thick Late Jurassic to Early Cretaceous of the Lower Saxony Basin in (a) and prominent basement faults associated with the Harz and Flechtingen highs in (b). Section (a) is compiled and simplified from Baldschuhn et al. (1996) and Mazur et al. (2005). Section (b) is from unpublished work by F. Jähne

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NGB and of specific problems they entail for strain analysis. We then discuss these concepts on the basis of several case studies that encompass the entire spectrum from normal fault reactivation and graben inversion to large-scale basement reverse faulting.

Figure 3.3.7. Balanced geometric models of a stratified succession (a), folded by horizontal shortening (b) or vertical salt uplift (c). Bed thickness is conserved in the first case whereas the limbs are thinned in the second. Although the magnitude of stretching as indicated by the strain ellipse is considerable in (c), limb thinning is not too conspicuous and may be difficult to detect or prove in real examples

Figure 3.3.8. Two line-length balanced models of end-member basement configurations, with equal amounts of shortening in the mechanically decoupled cover and basement. a) Undeformed succession. b) Pervasive basement deformation by relatively minor faults. Seismic control is good enough to preclude this as the sole mechanism for making strain in the basement and cover compatible. c) Single large basement fault feeds displacement into the salt detachment (Thin-skinned tectonics). This mechanism allows the basin floor to remain undeformed away from the master fault

Structural analysis and attempts to quantify the magnitude of strain in the NGB are faced with several challenges. First of all, as in any inverted basin, the structures observed today formed through polyphase deformation with at least one phase of extension and one of subsequent contraction. The recorded fault offsets are often only net displacements, i.e., the original normal offset minus the reverse offset produced by fault reactivation. The problem of separating the two components is much more severe on the basin margins where syntectonic strata have been eroded, than in the basin interior where seismic data often provide sufficient control on fault movements over time. Even so, analogue modelling suggests that the effects of mild inversion can be easily overlooked and the magnitude of contraction underestimated (Eisenstadt and Withjack 1995). Other problems are related to the existence of the thick Zechstein salt layer (Kukla et al. this volume). In particular, the following effects of salt must be taken into account: 1. Purely salt tectonic (halokinetic) movements are difficult to separate from the effects of regional deformation. Since salt is able to move in both contractional and extensional settings, it can affect estimates of both these strain components. Upwarping due to salt diapirism may be mistaken for folding caused by horizontal contraction or vice versa. Kossow et al. (2000), for example, argued that contrary to earlier held views, regional contraction was significant for the growth of the salt pillows in the NGB. Theoretically, diapiric rise in the absence of contraction must stretch and thin the limbs of salt structures, providing a diagnostic criterion (Fig. 3.3.7). Yet, this effect is quite subdued for salt pillows (Fig. 3.3.7c) and it may be difficult to detect it, especially on seismic time sections before depth conversion. Furthermore, stretching of the limbs will be accommodated by small faults rather than penetrative strain under the conditions of sedimentary basins, making it difficult to separate this strain from regional tectonic extension. 2. Regional sub-horizontal detachments can separate the sites of cover and basement deformation. Both in the LSB and in the NGB there is a marked discrepancy between the strongly structured sedimentary cover and the nearly flat or gently undulating basement surface (Fig. 3.3.6a, northern part of Fig. 3.3.6b, Fig. 3.3.10). The thick salt has obviously decoupled the deformation in the sedimentary cover from that in the

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Figure 3.3.9. (a) Depth map of base Zechstein (Permian) horizon showing main structural elements: N- to NNW-striking Permian normal faults, and NW-striking Allertal fault zone. Black arrows indicate the approximate kinematic direction during Permian extension. White dotted faults (labelled A, D, and G) belong to the Allertal fault zone, and are the same as those in figure 3.3.10. Grey areas are uninterpreted parts of the horizon due to overlying Zechstein salt diapirs. (b) 3D model illustrating a Permian normal fault with two Permian horizons and a seismic line. For location see grey box in (a). Horizon base Zechstein is colour-coded from red to blue with increasing depth

basement, but this does not explain why the strain magnitudes appear so different. The difference cannot be due to the basement being stronger than the overlying sediments, even though this may appear reasonable in terms of dynamics. Kinematic principles demand that basement and cover be strained by equal amounts. To illustrate the point, imagine a layer of sand on a wood-en board pinched in a vice. In spite of being easily deformable, the sand layer cannot shorten unless the board also does so. In order to make strain in the cover and basement compatible, two basic structural configurations are possible (Fig. 3.3.8). One configuration is to accumulate all fault offsets on a regional detachment linked to a basement master fault (“thin-skinned tectonics”). Though more common in contractional environments, this tectonic style has also been found in extensional settings and reproduced in analogue models (Brun and Nalpas 1996). The sites of mechanicallylinked basement and cover deformation can be separated by tens of km. A large structure (graben or anticline, respectively) with substantial basement relief of several km must evolve where the master fault penetrates the basement, most likely on the basin margin. With the second configuration, most basement deformation occurs within the basin, but is not apparent for one of two reasons: either the bulk strain could be distributed among a large number of small basement faults, or relatively large basement faults could be masked by reactivation that has restored the basement surface to a nearly flat geometry (near-zero net fault displacement). In the second case, the location of the large faults should be revealed by overlying grabens in the sedimentary cover. We will show that all these possibilities are realised to some degree in different parts of the NGB.

3.3.3.1 Reactivation of extensional structures: detailed 3D study around the western Allertal fault zone The western part of the Allertal fault zone separates the Lower Saxony Basin (LSB) and the Pompeckj Block (PB), as is indicated by differences in sedimentation (Best 1996; Betz et al. 1987; Hoffmann et al. 2001; Kockel 2003; Mazur and Scheck-Wenderoth 2005; Scheck-Wenderoth and Lamarche 2005), a high occurrence of Zechstein salt structures (Baldschuhn et al. 1996), and both vertical and horizontal movements (Stackebrandt and Franzke 1989; Betz et al. 1987). It has been described as a deep-seated Palaeozoic fault zone which underwent polyphase reactivation during the Mesozoic and part of the Cenozoic (e.g., Gast 1988; Brink et al. 1992; Hoffmann et al. 2005). These studies have shown that the deformation along the Allertal fault zone and within the LSB and the PB is very heterogeneous but the deformation style and the precise localisation of the Allertal fault zone, as well as the correlation between basement structures and Mesozoic structures are not well determined. Because these studies are mainly based on 1D or 2D investigations, they often cannot resolve complex 3D structures. When studying subsurface structures in 3D in more detail, the tectonic evolution appears more complex than observed in 2D seismic lines. In this section the heterogeneous deformation around a part of the Allertal fault zone is explained and illustrated in 3D, based on the interpretation of a 3D seismic data set with corresponding borehole data. For a more detailed description and discussion of this area see Lohr et al. (2007b).

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Figure 3.3.10. Thickness maps of Jurassic (a) and Late Cretaceous (b) sediments showing the main structural features and kinematics along the Allertal fault zone on a scale of several kilometres. White arrows indicate local kinematic directions. (c, d) Interpreted seismic crosssections through the Allertal fault zone. Sinuous line: Jurassic/Cretaceous unconformity, black arrows: normal faulting, white arrows: thrust faulting (inversion), white circles: strike-slip faulting with undefined shear sense. 3D seismic interpretation suggests that due to the high salt thickness in the western part (c), deformation is decoupled and partitioned in strike-slip kinematics below the Zechstein salt and dip-slip kinematics (extension, inversion) above the salt. Due to the low salt thickness in the eastern part (d), deformation between pre- and post-salt units is coupled and identified as strike-slip and oblique-slip kinematics

In the investigation area, the Allertal fault zone is observable as a NW-trending zone affecting Palaeozoic and Mesozoic rocks (Figs. 3.3.9, 3.3.10). The base Zechstein depth map (Fig. 3.3.9) shows major normal faults that developed during extensional deformation within the Southern Permian Basin. These faults grew synsedimentarily during the Permian, and their surfaces are strongly undulated, as shown in the 3D model in figure 3.3.9. From fault orientations, the kinematic direction during normal faulting is defined as roughly E-W extension (Fig. 3.3.9). These N to NNW trending Permian faults are truncated by the NW-trending Allertal fault zone (Fig. 3.3.9). Due to activity of this fault zone during the Mesozoic, some Permian normal faults show a present day offset of several hundred metres (e.g., along fault A, D, and G in Figs. 3.3.9, 3.3.10). As seen in the cross sections (Figs. 3.3.10c, d), an increased thickness of Keuper and Jurassic sediments demonstrates extension and graben development during that time. The fault pattern in the Jurassic thickness map (Fig. 3.3.10a) illustrates W to NW trending faults and grabens, from which the kinematic pattern is interpreted as NNE-

SSW extension with dip-slip normal faulting within postsalt units in the western part, and E-W extension associated with dextral oblique- to strike-slip faulting within both pre- and post-salt units in the eastern part (Fig. 3.3.11). Deformation in the Late Jurassic/Early Cretaceous was not observed along individual faults, but across the Allertal fault zone the Mesozoic sedimentary style differs strongly (Fig. 3.3.10). North of the fault zone Keuper sediments are condensed and Jurassic sediments are even missing, but Early Cretaceous sediments are thicker. South of the Allertal fault zone there is an increased thickness of Keuper and Jurassic sediments (Fig. 3.3.10a). Early Cretaceous strata overlie Dogger with an angular unconformity in the south, but rest disconformably on Keuper in the north. Due to the different levels of erosion, the Jurassic strata have a thickness of up to 1500 m in the south, but are absent in the north. The Late Cretaceous thickness map (Fig. 3.3.10b) shows no differences between the northern and the southern part of the study area. Areas of locally increased Late Cretaceous thickness represent salt rim synclines, whereas areas of locally reduced thickness are related to diapir

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this area, because changes in Zechstein salt thickness are of the same order (Fig. 3.3.10). Mesozoic faulting was followed by Cenozoic subsidence that is indicated by a widespread cover of Tertiary and Quaternary sediments of which the thickness increases towards the N, and is characterised by a general tectonic quiescence. The heterogeneous fault pattern associated with Late Cretaceous inversion indicates that deformation varies temporally from the Coniacian up to the Palaeocene along the Allertal fault zone (Fig. 3.3.10). This heterogeneity in distribution and timing of deformation is controlled by differential reactivation of older structures during extension, inversion and strike-slip faulting, and a variable salt distribution, which leads to decoupling of deformation between pre- and post-Zechstein rocks depending on the salt thickness. The development of salt structures is closely related to faulting along the Allertal fault zone and determines coupling (during strike-slip tectonics) or decoupling (during extensional and compressional tectonics) between pre- and post-salt units. Such heterogeneity points to stress perturbations and strain partitioning observable on different scales, not only along the Allertal fault zone but also within the Central European Basin System in general. Figure 3.3.11. Summary of tectonic events for the western Allertal area, with major structures and derived kinematic regimes. White arrows indicate horizontal extension direction; black arrows indicate horizontal compression direction: Z – Zechstein, BS – Buntsandstein, MK – Muschelkalk, Q – Quaternary

growth or to tectonic activity. Compressional faulting during the Late Cretaceous led to inversion of the Keuper and Jurassic grabens along the Allertal fault zone. In the western part, local N-W compression led to inversion which was concentrated predominantly along the former main graben fault E, which soled out into the Zechstein salt layer, and listric normal faults were reactivated as imbricate thrust faults (Fig. 3.3.10c). The vertical displacement along the imbricate thrusts is several hundred metres, but bedding-parallel shortening along the former main graben fault E is of the order of several kilometres. Correlation of tectonically controlled onlap structures with well data indicates that inversion occurred after the Coniacian and continued until the Maastrichtian, partly continuing until the Palaeocene. In the eastern part of the study area local E-W compression led to strike-slip faulting and inversion of a Jurassic pull-apart basin during the Santonian to Campanian. Further to the east of fault A onlap structures indicate uplift during Santonian to Maastrichtian of about 400 to 600 m, but this uplift was caused by salt migration underneath

3.3.3.2 Large basement reverse faults and associated thin-skinned thrusting: the Flechtingen High and Harz Mountains The Flechtingen basement uplift and the associated Gardelegen and Haldensleben faults (Fig. 3.3.6b) are part of the NEGB’s southern border. They are often regarded as a segment of the southern Elbe Fault zone (Elbe lineament, Franke and Hoffmann 1999a,b). Toward the SE, the Gardelegen and Haldensleben faults merge with the Wittenberg fault. As a first approximation, this system of folded and faulted basement forms a northeast vergent flexure whose front limb dips at varying angles and is broken by reverse faults to different degrees (Fig. 3.3.12). There is no clear evidence that older basement faults were reactivated during basin inversion in the foreland of the Gardelegen fault. North of the Gardelegen flexure, the deep reflection seismic line DEKORP Basin 9601 (DEKORP-BASIN Research Group 1999; Krawczyk et al. this volume) shows a nearly horizontal basement surface beneath the Zechstein salt overlain by complex structures in the suprasalt sedimentary cover. In Jurassic to Early Cretaceous times, the Zechstein salt in the southernmost NEGB accumulated into large salt pillows (Kossow et al. 2000). During Late Cretaceous inversion, these salt pillows were shortened to form three oblong, SE – NW trending salt anticlines which are in part strongly faulted.

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Figure 3.3.12. Block model of the Gardelegen basement flexure and its foreland. The sedimentary cover is only shown in two cross-section planes. The dashed outline shows the plane of the DEKORP basin 9601 seismic line, with the sedimentary cover shown in the south. In the east, the faulted transition zone from the Wittenberg fault to the Lausitz Escarpment forms the northwest dipping basin border. North of the Gardelegen and Wittenberg faults the base Zechstein dips gently N to NW. In the west (foreground) it dips W into the Gifhorn fault array. The crosssections give an impression of the complex deformation in the Mesozoic cover. South of the Wittenberg fault, erosional remnants of Mesozoic strata occur near major faults

This contractional deformation in the northern foreland of the Gardelegen uplift produced a maximum shortening of 2.5 km in the sedimentary cover. Yet, there is no indication of large basement thrust faults beneath the faulted salt anticlines that could have accommodated an equivalent amount of basement shortening. A thin–skinned model for the foreland structures with the Zechstein salt detachment linking them to a low-angle thrust fault underlying the Gardelegen basement flexure makes shortening in the cover and basement compatible. Folding in the basement was probably caused by the propagating master fault before it broke through onto the detachment. With these assumptions for a structural model constrained by the magnitudes of vertical uplift, horizontal shortening and the overall geometry of the basement uplift, balancing of the Gardelegen fault results in a listric shape and a 23–24.5 km deep intracrustal detachment under the Sub-

hercynian Basin. This coincides very well with reflectors imaged by the DEKORP experiment. Regardless of the balancing method applied, the displacement of the main basement fault at depth never exceeds 5 km. The Haldensleben and Wittenberg structures are faulted flexures with smaller amplitude than the Gardelegen flexure. The steep frontal faults of these structures dominate the structural geometry. To the southeast along the Wittenberg fault, the shortening is often transferred onto smaller basement faults covered by Mesozoic units (Otto 2003). The depth to basement also decreases southeastward from 4000 m below sea level to less than 1000 m below sea level. Due to the rising basement, no sediments younger than Keuper are preserved in the southeast. Shortening in the northern foreland decreases from 2 km in the west to less than 1 km at the southeastern end of Wittenberg fault. In the

Strain and Stress

Figure 3.3.13. Palaeostress map of the trajectories (trend, plunge) of the minimal horizontal principal stress axis (σhmin) of the Mesozoic dilatation period, calculated using the methods of Turner (1953) and Angelier (1979). Solid arrows (red): stress axis derived from fault kinematics, open arrows (yellow): stress axis derived from feather joint sets and mineral fibre growth. Numbers at the arrows indicate the plunges of stress axes greater than 10°. Cover rocks: yellow. MCR – Mid German Crystalline Rise/Kyffhäuser. Total number of analysed outcrops is 99

west, the Flechtingen basement uplift and the structures in its northern foreland terminate abruptly against the Gifhorn fault array. The Harz Mountains (peak elevation 1,140 m) represent an uplifted and deeply eroded block of Palaeozoic rocks. The Harz northern boundary fault separates this block from the Mesozoic cover of the foreland, particularly of the strongly subsided Subhercynian Cretaceous Basin in the north. This fault zone was possibly generated as a WNWESE striking ductile-brittle Variscan dextral strike-slip shear zone (e.g., Stackebrandt and Franzke 1989). However, a definite proof of this is difficult to obtain because Late Cretaceous thrusting nearly erased the traces of earlier events. This holds also true for the Early Mesozoic history of the boundary fault, but observations from surrounding areas allow to constrain its kinematics. Uniform thickness and facies trends of Triassic deposits exclude activity of

the fault in that time interval. Instead, the north-southtrending Eichsfeld-Altmark swell developed extending across the area that is now occupied by the western Harz Mountains. During Jurassic and Early Cretaceous times, the Harz area became an integral part of a system of dominantly NW-SE striking brittle normal fault zones, which developed between the Tornquist Zone and the North Sea grabens. In the exposed Palaeozoic basement of the Harz these faults were activated between the Jurassic and the Cretaceous as normal or oblique normal faults (Franzke et al. 2007). The NW-SE striking normal faults are partly filled with hydrothermal vein mineralisations. Radiometric age dating of alteration rims around the veins (K-Ar-dating on illite) as well as U-Pb-and U-He-dating on ore minerals (galena, hematite) bracket the time of vein formation between 200-90 Ma (Early Jurassic-early Late Cretaceous, Haack and Lauterjung 1993; Hagedorn and Lippolt 1993). In the Mesozoic deposits of the Harz foreland, faults of the

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same orientation occur as normal fault sets, synsedimentary grabens and half grabens (mostly of Early Cretaceous age, with a peak in the Aptian), cutting the complete sedimentary sequence. Palaeostress estimates compiled from fault sets in Mesozoic deposits of the Harz foreland and from mineralised structures within the Harz block show a strict SW-NE-orientation for the axes of minor principal stress (σ-hmin/ σ3). Under such stress conditions, NW-SE striking faults were activated as normal faults (Fig. 3.3.13). Up to the Cenomanian/Turonian there is no evidence for the existence of an elevated Harz block. Uplift and simultaneous subsidence of the northern forelands (Subhercynian Cretaceous Basin) started in Turonian to Coniacian (Voigt et al. 2004, this volume), indicating the onset of the Late Cretaceous inversion pulse. This deformation stage resulted in the formation of a more than 2 km wide flexural zone in Mesozoic sediments in front of the uprising Harz block. The northern boundary fault induced a

synsedimentary fault-propagation fold in the early stages of thrusting, and disrupted the stretched front limb of this flexure in a late stage. Drilling has shown that the emergent boundary fault dips about 50° to the SSW down to about 1 km depth (Fig. 3.3.14). Fission-track dating on apatites from late Variscan intrusives in the Harz Mountains indicates an enormous magnitude of uplift (4-5 km) in a short time span during the Late Cretaceous (85-82 Ma, Thomson et al. 1997). This process was accompanied by simultaneous erosion of the cover rocks from the Harz block up to the end of the Cretaceous (Voigt et al. this volume; von Eynatten et al. 2007). Taking into account that a pile of about 2 km of syntectonic Coniacian to Campanian sediments was deposited directly in front of the growing boundary fault (Subhercynian Cretaceous Basin, Roll 1953; Voigt et al. 2006), the magnitude of vertical displacement is at least 7 km, perhaps even 10 km. In figure 3.3.14, only the minimum displacements are calculated with respect to the present day erosion level.

Figure 3.3.14. Cross section through the northern boundary fault zone of the Harz Mountains. Location in figure 3.3.15. The dip of the main fault is well constrained to a depth of 1 km by the Schöth 2/65 drill hole. The minimum horizontal shortening (4 km) and the minimal vertical fault throw (>3 km) for the boundary fault are indicated. The hypothetical deep structure includes intense disharmonic folding as well as fracturing of more rigid layers at the synclinal bend between the steep front limb and subhorizontal foreland strata. Late Permian Zechstein salt may have invaded the evaporitic level of the Upper Buntsandstein and formed salt wedges. Fluid flow, probably along one of the splay thrusts, led to enhanced cementation of Cretaceous sandstones forming the cliffs of the “Devil’s Wall”. The Subhercynian Basin is probably detached from its basement along the Zechstein salt and thrust several km towards the north (see text). This interpretation requires a different structural geometry at depth as shown in the inset. See also figure 3.3.6b

Strain and Stress Figure 3.3.15. Palaeostress map of the inversion stage. The compressive tectonic event belongs to the inversion tectonics of the Late Cretaceous. The palaeostress axes (σHmax = σ1) were determined by indicators from 41 outcrops, supported by measurements of horizontal stylolites, drag folds and other increments in mineral veins. Cross section A-B shows the structural situation of the inversion stage. The Fallstein dome is interpreted as a flexural bulge

It can be demonstrated that normal faults in the sedimentary pile of the Subhercynian Basin, including the youngest sediments of the early Campanian, were still affected by small NNE-directed thrust faults (Franzke et al. 2007). Of minor importance are SSW-directed backthrusts. Older normal fault sets in the boundary flexure zone were rotated and partly transformed into reverse faults. Inside the uplifted Harz block, thrust or reverse faults can be found which cut hydrothermal vein fillings of Mesozoic age. The structures in the sedimentary cover of the northern Harz foreland are quite variable, with contractional, extensional and inverted salt structures occuring in close proximity. The effects of contraction generally diminish with increasing distance from the boundary fault, though not in an entirely systematic fashion. Closest to the boundary fault, thrusting of the Harz block was accompanied by cover folding and backthrusting that created the Harli, Quedlinburg and Asse structures (Fig. 3.3.15). The Quedlinburg anticline represents an inverted half-graben with a reactivated normal fault at its southern margin. The shortening during the inversion stage considerably over-

compensated the former dilatation. It was caused by an inversion of palaeostress. The principal stress axes kept their orientations (around SW-NE), but σHmax of the inversion stage replaced the σhmin axes of the former dilatational period. The Asse structure is a narrow salt anticline whose crest is dissected by a steeply-dipping thrust fault (Fig. 3.3.16). Since their hanging-wall cut-offs are often eroded, the shortening accommodated by the Asse and similar faulted salt anticlines is difficult to quantify. Depending on the fault and fold geometries used in restoration models and assumptions about the size and shape of the salt structures prior to inversion, the amount of horizontal shortening varies from more than 1 km to 2 km or more. Insufficient data resolution precludes more precise estimates. Farther north, the very gentle upwarp of the basement beneath the Huy and Fallstein anticlines (Figs. 3.3.6b, 3.3.15) may represents a flexural bulge developed at some distance from the overthrusting Harz block (Voigt et al. 2004), or a broad anticline created by basement folding. In the overlying sedimentary cover, the upwarp has been amplified by salt migration. In an even more distal position,

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Figure 3.3.16. Sketch cross-section of the Asse anticline as an example of contractional salt tectonics in the northern Harz foreland. Location in figure 3.3.5

the Offleben–Oschersleben salt wall is a good example of a purely extensional, salt-related structure (Fig. 3.3.6b). Farthest away from the Harz boundary fault and already close to the Flechtingen uplift, the eastern Allertal fault zone is a prominent but incompletely understood example of an inverted extensional salt structure (Fig. 3.3.17). There are indications of early salt movement in an extensional regime from Upper Buntsandstein to Jurassic time (Best 1996; Röhling 1991), creating a half-graben and associated reactive salt diapirism. The entire structure was slightly shortened by Late Cretaceous inversion, as is suggested by the anticline along the southern border of the structure, which folded or tilted Late Triassic to Cretaceous sediments. A qualitative restoration of the Allertal structure reveals a severe misfit between the high-angle fault cut-off on the northern border and the low-angle hanging-wall cut-off in the same Triassic strata imaged in the seismic data. This suggests that the two cut-offs can-

not belong to the same normal fault as interpreted earlier (Best 1996). (Fig. 3.3.17). As with the Flechtingen uplift discussed above, not all suprasalt structures in the Harz foreland are underlain by related basement structures. This suggests that a ductile detachment in the Zechstein salt and thin-skinned deformation also play an important role here. If this is true, then basement rocks of the northern Harz front must thrust over the detachment in the Zechstein salt at depth (Fig. 3.3.14, inset), with an overall architecture similar to the Gardelegen faulted flexure. The Harz, the Flechtingen uplift and other smaller basement uplifts in the region that formed as a consequence of Late Cretaceous shortening have similar structural geometries. However, they accommodated different amounts of strain, e.g., shortening of the Harz is more than twice that of the Flechtingen uplift.

Figure 3.3.17. Sketch cross-section of the eastern Allertal fault zone on the backlimb of the Flechtingen basement uplift. Location in figure 3.3.5

Strain and Stress

Distributed faulting and inverted structures of the basement 2D and 3D seismic data from the LSB and locally also the NEGB, show the top basement surface offset by steeply dipping faults that do not penetrate above the Zechstein salt (cf. Kukla et al. this volume). These faults can be quite frequent, but their displacements are a few 100 m at most and typically less. Most of them are interpreted as normal faults, but the resolution of older seismic data beneath the salt was often insufficient to reliably determine the fault type (F. Kockel, oral communication). While these faults were often interpreted and sometimes demonstrated by stratigraphic well control to be of late Rotliegend age (Gast and Gundlach 2006), the ductile salt decoupling them from overlying strata often precludes pinpointing their ages. Naturally, the fault spacing and recognisable throws depend on the resolution of the original data and the scale on which it is represented. Table 3.3.1 shows an analysis of distributed basement faulting from different parts of the LSB as interpreted in this chapter (Fig. 3.3.10) and two other publications. Fault spacing in these data is often between 1 and 2 km. The total vertical offset per km section length varies from 10 m to 100 m. For the sake of simplicity we assume that the horizontal offset (the heave) is half the throw for all faults. This corresponds to a dip angle of 63°, mechanically reasonable for normal faults. We can then estimate the extension accommodated by the fault set, which ranges from <1 % to 5 %. Over a larger area the small-scale faults thus represent only a few percent of extensional strain. Although no reliable estimate yet exists of the extension that formed the LSB, the stratigraphic restoration of Adriasola-Muñoz et al. (2007, see also this volume, Fig. 3.4.16) permits an approximation of magnitudes of crustal thinning and extension. Depending on the densities chosen for mantle, crust and sediment infill, the crust must thin by 5-7 km to accommodate an average 3.5 km of Jurassic-Early Cretaceous strata. This corresponds to a 14 % to 25 % horizontal

extension for an initially 40 km or 35 km thick crust. So, even if all the small-scale faults were taken to have been active in Mesozoic time, they could at best provide one third of the total extension required to create the basin. A more substantial contribution to concealed basement extension and shortening must come from reactivated normal faults which now have low net displacements. Figure 3.3.18 shows part of a structural section across the inverted Rheder Moor fault zone on the northern margin of the LSB, taken from Baldschuhn et al. (1996). If the section is restored to a pre-inversion configuration with no change to the basement geometry, the hanging-wall of the main fault overlaps with the basement. The basement surface has to be lowered by some 1000 m to accommodate the hanging-wall strata, substantially increasing the amount of normal displacement and reverse offset on the underlying basement fault. A similar situation is depicted in the 3D seismic data of figure 3.3.10, where the base Cretaceous unconformity cannot be restored to horizontal unless the basement south of the Allertal fault zone is lowered by several hundred meters. Another case, but without recognizable basement faults, may be present in figure 3.3.6a, where a large accumulation of Jurassic to Early Cretaceous strata with a distinct “turtle” geometry occurs between the Steinhude and Allertal fault zones, above the area labelled “Blind basement faults?”. This structure is not primarily related to salt withdrawal (the diapir apparently bounding it in the SW actually lies some 30 km away on the other segment of the section). Rather, it must represent a former basement syncline whose trough was raised again during inversion to arrive at a nearly flat basement surface. Although the structure is smaller and the degree of inversion is less, its overall architecture is strongly reminiscent of the Pomeranian segment of the inverted Mid-Polish trough (Krzywiec 2002; Mazur et al. 2005).

Table 3.3.1. Geometrical parameters of distributed basement faults Source

Baldschuhn et al. 2001 This chapter, Figure 3.3.10c This chapter, Figure 3.3.10d Mohr et al. 2005, Figure 4 Mohr et al. 2005, Figure 5 Mohr et al. 2005, Figure 6

Profile /section Number of length (km) faults

Average fault spacing (km)

Sum of vertical Sum of horizontal Vertical offset per Extension offsets (km) offsets (km) at 63° km section length (%) fault dip (km)

10

8

1.25

1

0.5

0.1

5

13.3

7

1.9

0.28

0.14

0.02

1.05

17.1

3

5.7

0.25

0.125

0.01

0.73

29.3

17

1.7

2.1

1.05

0.07

3.58

13

9

1.4

1.24

0.62

0.10

4.77

13

10

1.3

1.06

0.53

0.08

4.08

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J. Kley et al. Figure 3.3.18. Evidence for large normal-sense fault offset reset to near zero by reverse reactivation at the northern border of the LSB. a) Part of a regional cross-section constrained by seismic and well data (Baldschuhn et al. 1996). b) Restoring the suprasalt succession to its pre-inversion geometry (base Late Cretaceous flattened out) with an unmodified basement surface results in an invalid overlap of hanging-wall strata and basement. c) Restoring one of the minor basement faults (red arrow) to a large normal offset produces a valid restoration. Most of this normal displacement must have been consumed during inversion. Location in figure 3.3.5

The examples discussed above show that the styles of basement and cover deformation and the way they are linked vary strongly across the NGB. Thin-skinned thrusting is well developed on the southern borders of the NGB and the Subhercynian Basin. In contrast, there is no clear evidence for regional detachments in the LSB. Cover deformation is more closely tied to underlying basement structures which are sometimes obscured by inversion, and distributed basement strain is more important.

3.3.4 Stress history 3.3.4.1 Palaeostress analysis from outcrop and seismic data Knowledge of the palaeostress states is essential to understand the tectonic evolution of the crust. Classical palaeostress methods (Angelier 1994) employ the analysis of

Strain and Stress

Box 3.3.3 Palaeostress measurements Palaeostress states can be inferred from deformation structures on different scales. On microstructural scale, calcite e-twin lamellae and quartz deformation lamellae are assumed to represent planes of high resolved shear stress which have a restricted range of orientations to the principal stress axes. On larger scales, faults with known sense of slip may provide the basis for the inversion of palaeostress states (fault-slip analysis).

Outcrop scale Conventional fault-slip analysis uses faults from outcrops, i.e., faults on a 10-1 to 102 m-scale, to derive the causative palaeostress state. A fault-slip datum consists of the fault plane orientation, the slip direction (striae like slickensides) and the sense of slip (reverse, normal, dextral, or sinistral); the latter can be estimated by kinematic indicators (Fig. 1; Doblas 1998). Crosscutting relationships between different faults or differently orientated striae on the same fault plane are expressions of successively acting stress states and the best tools to separate heterogeneous fault populations. If such chronological criteria are insufficiently available in the field, alternative computer-aided separation tools must be applied (Sperner et al. 1993; Nemcok and Lisle 1995; Yamaji 2000). All stress inversion techniques based on fault-slip data are either related to slip criteria or frictional criteria (Angelier 1994; Ramsay and Lisle 2000). According to Wallace (1951) and Bott (1959), one can predict the slip direction along any plane of known orientation under a given stress state assuming that slip takes place in the direction of maximum shear stress (WallaceBott hypothesis). Knowing the slip directions of many faults allows numerical determination of the “reduced stress tensor” (Angelier 1979), which consists of (i) the directions of the principal stress axes σ1≥σ2≥σ3 and (ii) the ratio of principal stress differences R=(σ2-σ3)/(σ1-σ3). Contrary to recent stress measurements, fault-slip analysis does not yield stress magnitudes. Other approaches assume each slip plane to be newly formed, i.e., fractured and slipped under the same stress state (Numeric Dynamic Analysis, PBT-Method; Sperner et al. 1993). In this case the orientation of the fault plane allows the relative magnitudes of the shear stress τ and the normal stress σn in the plane to meet with the Mohr-Coulomb yield criterion, τ=C+µσn (C is the cohesion, µ is the coefficient of friction). According to principles of fracture mechanics, a defined fracture angle θ (usually θ=30°) between each fault plane and its associated contraction axis ε3 is adopted to construct the strain axes ε3, ε2, and ε1 for each fault-slip datum. The bulk reduced stress tensor can be estimated by cluster analysis of the strain axes of the whole number of striated faults. Therefore parallelism between the average ε3 and σ1, between ε2 and σ2, and between ε1 and σ3 is assumed.

Seismic scale Fault-slip analysis on a seismic scale has its very specific challenges. Methods for determining slip direction from 3D seismic data include (i) displaced linear objects like channels or dikes (Back et al. 2006), (ii) position of hanging wall and footwall cutoffs and associated variable displacement distribution along fault strike (Allan 1989; McLeod et al. 2000), or (iii) corrugations on undulated fault surfaces caused by fault segment linkage (Needham et al. 1996; Marchal et al. 2003; Lohr et al. in press). Workflow for reconstructing movement directions: (i) detailed 3D interpretation of horizons and faults from depth-migrated 3D seismic data, (ii) analysing deformation history and determining slip direction of all incremental kinematics, (iii) subtracting younger increments from older ones by successive backward modelling, (iv) fault slip data inversion.

Figure 1. Calcite slickensides on a subvertical strike-slip fault within Late Carboniferous andesites from the Southern Inverted Margin area (Mammendorf quarry). The arrow indicates the derived dextral sense of slip

Figure 2. 3D model of an undulated fault surface interpreted from a depth-migrated 3D seismic data set. From the corrugations on the surface, kinematic movement directions can be derived. The orientation of the corrugation axis is used as fault surface azimuth/dip, and equally as bearing/plunge of the “striae”, whereas the sense of slip of the hanging wall is derived from the 3D seismic data. White arrows are examples of corrugation axes. No vertical exaggeration

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fault populations accessible in outcrops. In covered areas of sedimentary basins, as in large parts of the CEBS, highresolution 3D seismic data can be used to infer palaeostress states. Studies that investigate the structural evolution of the CEBS on a basin scale (Scheck-Wenderoth and Lamarche 2005) or on the scale of sub-basins (Scheck and Bayer 1999; Hansen et al. 2000; Baldschuhn et al. 2001; Scheck et al. 2002a,b; Lamarche et al. 2003) argue for a changing stress field affecting the CEBS over time. This is also confirmed by various palaeostress studies in other parts of north-central Europe (e.g., Bergerat 1987; Vandycke 2002; Lamarche et al. 2002; Bergerat et al. 2007). To elucidate the deformation and associated palaeostress history of a basin system, fault-slip analysis is an appropriate tool (Box 3.3.3). From a 3D seismic data set in the western Allertal fault zone (location in Fig. 3.3.5) numerous Permian normal faults have been interpreted in 3D and analysed with respect to their morphology. All faults show a strong undulation of their surfaces (Fig. 3.3.19). Such undulations usually develop as a result of propagation and coalescence of several smaller fault segments through time (McLeod et al. 2000; Marchal et al. 2003). Corrugations on these surfaces can be highlighted by applying attributes such as dip, azimuth, or Gaussian curvature (Lohr et al. in press). The fault corrugations should be orientated parallel to the slip direction because this movement should require the least energy and therefore the smallest strain, and thus corrugations indicate the movement direction of the hanging wall along the fault (Needham et al. 1996). Fault-slip analysis requires the orientations of fault planes, and their directions and senses of slip. From each corrugation (n=38) the azimuth and dip of all surface triangles was measured. Their average value refers to the orientation of corrugation axis, which was used as fault plane orientation, and equally as direction of slip. The sense of slip of the hanging wall was derived from the seismic data. Principal stress directions were calculated by using the method of Numeric Dynamic Analysis (Spang 1972), and the associated computer program NDA (Sperner et al. 1993). The Permian normal faults document a triaxial stress state (R=[σ2-σ3]/[σ1-σ3]=0.4) during the Rotliegend (Permian) with a vertical axis of maximum compression, σ1, and an ENE-WSW directed tension axis, σ3 (Figs. 3.3.20). From another 3D seismic data set in the Groningen area (location in Fig. 3.3.5) several normal faults truncating Upper Rotliegend (Permian) strata were analysed using fault surface geometries to derive kinematics (van Gent et al. in press). The slip direction and sense was determined by an-

alysing and comparing minima and maxima distance between corresponding footwall and hanging-wall horizon cut-offs, and to a lesser degree from surface corrugations. For the calculation of stress states NDA was applied to the fault-slip data sets. Regardless of the approach towards finding slip directions, the Rotliegend sediments document a triaxial stress state (R=0.4) with a vertical axis of maximum compression, σ1, and an ENE-WSW to NE-SW directed minimum compression axis, σ3 (Fig. 3.3.20). The stress direction derived for the Rotliegend fits very well with that observed in the Allertal area. Interpretation of the base Late Cretaceous indicates a period of tectonic quiescence at the beginning of the Late Cretaceous. The palaeostress results of deformation during the Late Cretaceous indicates a triaxial stress state (R=0.4) with a vertical axis of maximum compression, σ1, and a WNW-ESE directed σ3 axis (Fig. 3.3.20). This observation fits well with field studies presented by Vandycke (2002) for outcrops in Sussex and the Belgian Mons Basin. Unlike the southern margin of the CEBS (see below), the Groningen area was not inverted by Late Cretaceous / Early Tertiary tectonics (de Jager 2003). Four horizons were interpreted and reconstructed from the Cenozoic North Sea Super Group in the Groningen area. The analysis of two horizons at the base of this Super Group indicates a period of tectonic quiescence. Since this reconstruction proves the Cretaceous deformation in the Groningen area to be older than the Cretaceous-Tertiary boundary, the WNW-ESE extension is interpreted to predate the Late Cretaceous / Early Tertiary NE-SW compression observed at the southern margin of the CEBS. Combining the palaeostress results from two other Cenozoic horizons, a triaxial stress state (R=0.3) with a vertical axis of maximum compression, σ1, and an NW-SE directed σ2 axis can be derived (Fig. 3.3.20). This observation fits well with field studies presented by Vandycke (2002) for outcrops in Sussex, Kent, NE Belgium and the Belgian Mons Basin. Since the youngest of the interpreted horizons falls just within the neotectonic period (as defined by Becker 1993 for NW Europe), comparison with the present stress fields is justified. The Cenozoic results from this study fit well with the stress state described in the World Stress Map (WSM, Reinecker et al. 2005; compare Fig. 3.1.21) and well data from an unpublished study. Patterns of mesoscale faults (constituting more than 850 fault-slip data) from more than 40 outcrops within Late Carboniferous, Middle Triassic (Muschelkalk), Late Jurassic, and Late Cretaceous rocks provide the basis for palaeostress analysis along the southern inverted margin of the CEBS (Fig. 3.3.21, Sippel et al. in press). Fault-slip analysis was performed for each station by combining the P-B-T-Method (Turner 1953, after a so-

Strain and Stress Figure 3.3.19. Permian normal faults in map view colour-coded by depth, interpreted from a 3D seismic data set in the western Allertal area. Undulations on the fault surfaces are interpreted mainly as corrugations, which are used as kinematic vectors from which palaeostress data have been derived. The plane-lineation-movement sense data and the principal stress directions are plotted in a Schmidt net, lower hemisphere, equal area projection. Used method is NDA, numerical dynamic analysis

lution of Sperner et al. 1993) with the Multiple Inverse Method (MIM; Yamaji 2000) as described by Sippel et al. (in press). Many sites display a heterogeneous fault population accompanied by a considerable number of reactivated fault planes, both attesting polyphase deformation. One inferred stress state, however, is preserved by almost all sites regardless of the age of the host rock: it is characterised by a NNE-SSW- to NE-SW-directed maximum principal stress axis (σ1) and relatively low stress ratios (0.1≤R≤0.3; southern inverted margin area, Fig. 3.3.20). Its prevalence, as well as its high consistency in terms of both the direction of maximum compression and the stress ratio R, documents the high relevance of this stress state for the southern margin of the CEBS. The permuting directions of the minor principal stress axes (σ2 and σ3), in contrast, seem to be related to local variations of the stress field. As the youngest rocks preserving this NE-SW-compression are of Turonian age, this stress configuration must be assigned to Late Cretaceous or later times when large parts of the basin system were tectonically inverted due to the far-field compressional stresses induced by Africa-Europe convergence. The character and timing of the NE-SW-compression relate it to the well-constrained Santonian-Campanian inversion pulse in the NGB (e.g., Ziegler 1987; Kockel 2003; Voigt et al. 2004) and to the Maastrichtian-Palaeocene compression that caused the inversion of the Polish Trough (Lamarche et al. 2002). The imprints of other stress states derived from outcrops along the southern margin of the CEBS, whether older or younger than the NE-SW-compression, are much weaker and limited to much smaller sub-areas.

3.3.4.2 Neotectonics, seismicity and present-day stress state Lithospheric large-scale folding, intracratonic faults and the recognition of the causal reactivation source play a crucial role in understanding the neotectonics and the post-glacial landscape evolution of the CEBS (Cloetingh et al. 2004). Major basement faults are directed NW-SE, minor faults NE-SW and NNE-SSW (Figs. 3.3.4, 3.3.5). The major tectonic stress field currently acting within large parts of the CEBS is induced by the ongoing Alpine collision, and is directed NW-SE to N-S (Grünthal and Stromeyer 1992; Reinecker et al. 2005), being almost stable since the mid-Tertiary (Oligocene). The CEBS was partly affected by glacial loading and unloading of the ice masses during the Pleistocene. The postglacial uplift and the re-adjustment of the lithosphere significantly influence the regional stress field (see Sirocko et al. this volume). Present-day stresses in the North German Basin, mainly inferred from earthquake focal mechanisms (Gregersen 1992; Müller et al. 1992), have been compiled recently by Reicherter et al. (2005). (See Box 3.3.4 for a brief description of techniques used in the analysis of present-day stress). The effects of plate stresses on the tectonics of northwestern Europe cannot be sufficiently explained by Atlantic ridge push forces and the Alpine convergence (Goes et al. 2000b). Deviation of the recent stress in subsalt layers differs from the general West European trend, which is directed NW-SE (compare Figs. 3.1.14, 3.1.21). Roth et al. (1999) found a N to NE directed SHmax, mainly

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Figure 3.3.20. Overview map showing results of palaeostress analyses for the study areas Groningen, western Allertal, and southern inverted margin of the CEBS

in the area east of Hamburg in Mecklenburg-Vorpommern, based on borehole breakout data. In contrast, maximum horizontal stresses in supra-salt layers point to W-directed compression. The regional mantle structure is thought to play an important role in the intraplate tectonic activity in Central Europe and may act as the major agent in the post-glacial landscape evolution by differential decompensation of the ice loading/unloading (Reicherter et al. 2005). The dominant influence of the post-glacial rebound on the stress field in northern Germany has been shown by GPS data (Scherneck et al. 1998) and finite element modelling of the stress field (Kaiser et al. 2005; Cacace et al. this volume). Strong lithospheric barriers lead to significant deviation of horizontal stresses. The STZ and the rigid Bohemian Massif are such barriers which transfer stress into their forelands. However, the regional and local impacts of the complex interplay between glaciation/deglaciation and deep-rooted geodynamics yield evidence of the lithospheric memory effect on the landscape of the CEBS (Reicherter et al. 2005).

The Scandinavian post-glacial rebound dome and glacial forebulge in relation to the underlying tectonic stress field have engendered a pattern of alternating regions of higher and lower seismicity. The most dramatic evidence of active faulting and seismicity is reported from those areas where ice sheets reached their maximum thickness, including the marginal regions (Stewart et al. 2000; Mörner 2003). One effect is relative sea level changes in the Baltic Sea and North Sea areas caused by ice melting and a still rising rebound dome (Mörner 1973; Behre 2003). The recovery from Late Pleistocene glaciation by elastic lithospheric flexuring and lateral mantle flows affects large areas (Wu et al. 1999; Thorson 2000). The area in Northern Germany can be regarded as a marginal area of recovery, but also experienced deglaciation from an approximately 1000 m thick ice sheet (Piotrowski and Tulaczyk 1999) at the end of the Pleistocene. Measurable crustal deformation is still the consequence of the mantle response to deglaciation (Scherneck et al. 1998), accompanied by decelerating seismic activity (e.g., earthquake

Strain and Stress

Box 3.3.4 Recent stress measurements Borehole breakouts are stress-induced failures of the borehole, often caused by conjugate fractures and material spalling. Breakouts are parallel to minimum horizontal principal stress axis, and their geometry is controlled by rock strength, lithology, pre-existing fractures, drilling intensity and method, and borehole diameter (Gough and Bell 1982; Zoback et al. 1985; Plumb and Hickman 1985). DITF (drilling induced tensile fractures) are tensile fractures that form parallel to maximum horizontal principal stress. They are measured by downhole sonic televiewers or springloaded calliper tools. DITF are used for measuring orientations and magnitudes of principal stress in boreholes where conventional hydraulic fracturing is difficult, as it provides a continuous stress profile along the borehole (Moos and Zoback 1990; Peška and Zoback 1995; Brudy and Zoback 1999). Overcoring is a stress relief technique assuming that stress on an elastic material produces a proportional strain. Overcoring involves drilling an annulus around a hole to form a hollow cylinder of rock on which the stress from the surrounding rock has been released. The stress release causes elastic deformation of the cylinder, which is used to calculate magnitude and direction of stress (Twiss and Moores 1992; Ljunggren et al. 2003; Sjöberg et al. 2003) With Hydrofracturing the differential stress (Dσ= σ1-σ3) is measured in-situ to determine magnitude and orientations of principal stress. Fluid is pumped up in a borehole section until a tension fracture forms at critical pressure Pc. Fluid is sealed into the fracture immediately, and pressure drops and stabilises at a value that is just sufficient to keep the fracture open. Hydrofractures form parallel to maximum horizontal stress (Brudy and Zoback 1999; Vernik and Zoback 1992; Bjarnason et al. 1989; Ljunggren et al. 2003) In seismic regions, earthquakes are the result of sudden energy release due to stresses in the Earth. The first motions of P waves (pressure, primary) and location of aftershocks indicate shear sense and fault orientation, and used to construct focal mechanisms, so-called “beach balls”. Main compressional stress axes lie within P dihedra, minimum stress axes lie in tensile dihedral. (Shearer 1999; Mussett and Khan 2000). High resolution earth observations such as Global Positioning System (GPS) and Interferometric Synthetic Aperture Radar (InSAR) using satellites enable precise measurements of horizontal and vertical surface movements, from which stress data can be derived (Massonnet et al. 1993; Massonnet and Feigl 1998; Hanssen 2001). Interested readers are especially referred to Ljunggren et al. (2003) and references therein, in which the most common methods are briefly described including application areas and limiting factors. Results of recent stress measurements are summarised in the World Stress Map: http://www-wsm.physik.uni-karlsruhe.de/pub/introduction/introduction_frame.html

catalogues of Sweden, Mörner 2003; and the German BGR, Hannover). The seismicity in northern Germany is rather low; the probability of earthquakes larger than MSK intensity > 5, as documented in the earthquake catalogue for the last 1000 years (Leydecker 1986), is also very small and on the order of 10,000 years (Leydecker et al. 1999). Rare earthquakes and fault plane solutions indicate N-NW directed maximum horizontal compression (Reicherter et al. 2005).

3.3.5 The CEBS´s structural evolution From latest Carboniferous time to the present, the structures in the CEBS evolved in 4 major stages corresponding to changing kinematic (strain) and dynamic (stress) regimes. The varying strain patterns during these intervals are relatively well constrained, but particularly in Cenozoic time are too complex to be explained by a single, consistent, far-field plate tectonic stress regime

(Cacace et al. this volume). This indicates that stresses of more local origin must have played a role in the CEBS’s structural evolution. The four stages are briefly described below, accompanied by interpretative sketch maps of active structures and kinematic regimes (Fig. 3.3.22). The Mesozoic extension stage is divided into two substages. Stage 1 (Fig. 3.3.22a): Transtension in latest Carboniferous to Permian time (Rotliegend). Widespread voluminous volcanism is accompanied and followed by the evolution of normal and transtensional faults with associated sedimentary basins. This fault system has been interpreted to be part of a wide E-W trending belt of dextral transcurrent motion connecting the still-active Ural and Appalachian thrust belts in the latest phase of the Variscan orogeny (Arthaud and Matte 1977). At any rate, the area affected extends beyond the future CEBS whose principal features became established as faulting waned and thermal subsidence took over in late Rotliegend to Zechstein time (van Wees et al. 2000).

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Stage 2a (Fig. 3.3.22 b): Extension in Early to Late Triassic (Middle Jurassic) time. Substantial E-W extension on N-S trending normal faults produces the North Sea Central, Horn and Glückstadt Graben, as well as the Weser and Rheinsberg troughs. NW-trending faults partly linked the N-trending grabens. Normal faulting began in Buntsandstein times (Röhling 1991), but the main extension phase is considered to be of Keuper age. This is also the time when many salt structures begin to evolve. Extension of several km has been interpreted for individual, salt-intruded fault zones of both N and NW trend (Best 1996). Coeval, substantial NE-SW extension or dextral transtension occurs along the Tornquist zone, suggesting some degree of strain partitioning in a general E-W to NE-SW extension regime. Little faulting is documented in the area south of the large grabens. Stage 2b (Fig. 3.3.22c): Extension in Late Jurassic to early Late Cretaceous time. Strong extension creates the Lower Saxony Basin south of the Rheder Moor and Allertal fault zones and west of the Gifhorn system. The most important normal fault system along the southern border of the LSB creates as much as 7 km of vertical offset. The NEGB has no large normal faults during this time interval. The switch from the dominantly N-S trend of the Triassic grabens to the NW-SE structural trend of the LSB suggests an anticlockwise rotation of the extension direction. Alternatively, or additionally, the NW-SE trending normal faults may have had substantial dextral strike-slip components. The reverse basement faults now bounding the NEGB in the south are sometimes interpreted to have been active during this period, but their SW dips preclude interpreting them as original syndepositional, down-

to-the-basin normal faults. Regional uplift involves the northerly adjacent area that had been affected by Triassic E-W extension, creating the Pompeckj Block along the southern flank of the Ringkøbing-Fyn high. Stage 3 (Fig. 3.3.22d): Contraction and inversion in latest Cretaceous to Late Oligocene time. Extension rapidly gives way to contraction. The main extensional systems of stage 2b become reactivated as reverse faults. The LSB as well as the Norwegian-Danish and Polish basins along the Tornquist Line are partially inverted. Basement reverse faults, probably newly formed, dissect and raise the southern margin of the NEGB. The general contraction direction at the onset of this stage is SSW-NNE, although strong local deviations have been documented (e.g., Fig. 3.3.11). Despite frequent claims to the contrary (Brun and Nalpas 1996; Deeks and Thomas 1995; Krzywiec et al. 2003; Wrede 1988), no unambiguous evidence seems to exist for a regionally consistent strike-slip component on the NW-trending faults. This is borne out by the fact that both sinistral (Betz et al. 1987; Ziegler 1990) and, more frequently, dextral strike-slip components have been proposed (references above). However, there is ample field evidence for small-scale conjugate strike-slip faulting consistent with a NNE to NE compression direction (Sippel in press; section 3.3.4 and Box 3.3.3). Mostly milder, N-S to NE-SW contraction and inversion continue into the Palaeogene in the southern North Sea and the Polish basin, although it has been argued that uplift in Palaeocene time is due to elastic rebound of Late Cretaceous basins upon relaxation of the regional stress (Nielsen et al. 2005). The late phases of inversion in Eocene time are coeval with W-E to NW-SE directed extension creating

Strain and Stress

Figure 3.3.22. Main structural evolution stages in the CEBS. These maps are intended to show the overall pattern and extent of fault activity in the different time slices. They should not be taken as very reliable indicators of the activity or inactivity of individual faults, in particular not the minor ones

the Upper Rhine and Eger rifts as parts of the European Cenozoic Rift System (Dèzes et al. 2004). Although the major horizontal stress direction is the same for the inversion structures and the grabens, the dynamics of coexisting inversion and rifting are not yet fully understood

(Dèzes et al. 2005; Michon and Merle 2005). In the NGB, regional gentle tilting toward an area of strong subsidence in the southern North Sea (Hall and White 1994) is superimposed on mostly minor faulting sometimes associated with salt movements. However, more substantial

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extensional reactivation of some faults is also observed in the CEBS, e.g., in the Glückstadt graben (Kockel 2003; Maystrenko et al. 2005a,b). Stage 4 (Fig. 3.3.23e): Complex kinematics in Neogene time. Some time between the Late Eocene and Middle Miocene, the major horizontal stress rotated anticlockwise from a NE-SW to a NW-SE direction, attaining its present-day orientation (Bergerat 1987; Hinzen 2003, see section 3.3.4.2). This is well documented by NW-direct-

ed thrusting in Switzerland which creates the Jura foldand-thrust belt and raises the external basement massifs. Away from the Alpine front, however, this event is not prominently expressed. One set of horizontal stylolites indicates NW-SE compression (Kurze and Necke 1979) and, since it continues into the Jura fold belt, may be of Neogene age. NE-trending folds and faults occur locally, but the ages of these structures are not well constrained. As in Palaeogene time, the Neogene contraction coexists with extension in the Upper Rhine, Lower Rhine and Eger

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Chapter 3.4

Subsidence, inversion and evolution of the thermal field R. Littke · M. Scheck-Wenderoth · M.R. Brix · S. Nelskamp

3.4.1 Introduction Large sedimentary basins such as the Central European Basin system (CEBS) can be regarded as geo-reactors, in which many important chemical reactions take place, leading e.g., to the formation of petroleum, natural gas and coal. These reactions are kinetically controlled. Understanding of temperature evolution over time is therefore of utmost importance in basin studies in general and in petroleum exploration in particular. Temperature is closely coupled to crustal heat flow, radiogenic heat production, convection of fluids, and – most importantly - to depth of burial of sedimentary rocks. For this reason, analysis of subsidence and inversion in the framework of numerical burial history reconstructions has become a widely applied tool in basin research. This chapter will first present the present-day thickness/depth pattern of the sediment fill of the CEBS as well as the present-day temperature field in the context of some basic information on temperature in sediments. Thereafter, maturity and other temperature-sensitive parameters are presented which are commonly used to deduce past temperatures and burial depths of sedimentary rocks and serve as calibration parameters in numerical basin modelling. Finally, the application of these techniques to the CEBS is described, mainly based on several cross sections through the basin, for which temperature evolution over time is shown in the context of subsidence and inversion.

3.4.2 The CEBS as example of regional subsidence models To understand the controlling factors of present and past processes in intra-continental basins such as the CEBS, it is important to assess the spatial interplay of uplift, subsidence, and related mechanisms, including deep-seated processes. Therefore, the consistent integration of all available geological and geophysical data is required to describe the present day physical state of basins and to reconstruct basin history. A range of geological and geophysical methods such as interpretation of well, seismic and seismologic data, structural analysis and restoration as well as modelling of the thermal and potential fields

all help to unravel the physical processes controlling basin evolution and the related changes in temperature and pressure conditions. Both the changing regional pattern of subsidence and uplift over time and the assessment of the internal deformation, as for example related to salt mobilisation, need to be studied to find the main boundary conditions determining the evolution of the geo-reactor. In the CEBS, the regional pattern of subsidence changed considerably over time and structural analysis of these changes reveals different causal mechanisms of individual tectonic regimes. Subsequent to the Variscan orogeny, the lithosphere of Central Europe was subjected to a series of tectonic events which were related to the ongoing break-up of Pangaea. In response to these processes, the tectonic evolution of the CEBS was determined by its position between the stable Precambrian Baltic-East European craton in the N and NW and the competing mega-rift systems of the Arctic-North-Atlantic in the NW, the central Atlantic in the W, and the Tethyan oceans in the S. During the later Mesozoic and Early Cenozoic, ongoing rifting resulted in crustal separation in the North Atlantic, whereas the successive closure of the Tethyan oceanic basins and continental collision between Africa and Eurasia caused compression and basin inversion in the CEBS. The resulting superposition of plate-boundary-induced stresses led to the development of a complex structural pattern in the CEBS with subsidence and subsequent inversion of numerous sub-basins. The CEBS contains the thickest Permian-Cenozoic succession (>10 km) in central Europe (Scheck-Wenderoth and Lamarche 2005) and the base of this sedimentary fill, the top pre-Permian surface (Fig. 3.4.1), illustrates the different sub-basins which have developed since latest Carboniferous times. Subsidence and uplift within the CEBS were strongly influenced by two major groups of structures, oriented NW-SE and N-S. This is obvious at the top Pre-Permian surface. The NW-SE trend controls the shape of three large basins, namely: (i) the Polish Basin, (ii), the Norwegian-Danish Basin and (iii) the North German Basin extending westward into the southern North Sea. NW-SE directed structural control is furthermore provided by the Sorgenfrei-Tornquist Zone (STZ) at the northern

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Figure 3.4.1. Structural overview of the Central European Basin System depth to top-pre-Permian surface with major NW-SE-oriented fault systems: (i) the Tornquist-Zone consisting of the Sorgenfrei-Tornquist-Zone (STZ) north of the Norwegian-Danish Basin and the TeisseyreTornquist Zone (TTZ) beneath the axial part of the Polish Basin and (ii) the Elbe Fault System at the southern margin of the North German Basin. The large Permo-Mesozoic Norwegian-Danish and North German basins are separated by the Ringkøbing-Fyn-High. Between the NW-SE striking fault systems, N-S-directed structural elements are present, the largest of which are the Central Graben (CG), the Horn Graben (HG) the Glückstadt Graben (GG) and the Rheinsberg Trough (RT)

margin of the Norwegian-Danish Basin, by the TeisseyreTornquist Zone (TTZ) along the axis of the Polish Basin, by the Mid North Sea-Ringkøbing-Fyn-High (RFH) chain of basement highs that separates the Norwegian-Danish and the North German basins (Cartwright 1987) and by the Elbe Fault System (EFS), bounding the North German Basin to the S (Scheck et al. 2002a; Maystrenko et al. this volume). The second group of roughly N-S- trending structures is represented by large grabens such as the Central Graben, the Horn Graben, the Glückstadt Graben and the Rheinsberg Trough (Fig. 3.4.1). When each of the main structural elements was active is best illustrated by the shifts in depositional areas over time. As a first order approximation, the thickness distribution of the different stratigraphic units provides an overview of the regional pattern.

3.4.2.1 Late Carboniferous-Early Permian The oldest unit of the post-Variscan succession consists of Permo-Carboniferous volcanic rocks (Breitkreuz et al. this volume) which are the traces of basin initiation. Though the related mechanisms are still debated, regional thermal destabilisation and a transtensional palaeo-stress field

(Knott et al. 1993; Ziegler 1990; Arthaud and Matte 1977) appear to have been the key controlling factors prior to Permian subsidence (Ziegler 1990). This resulted in magmatic activity and extensive volcanism that affected nearly the whole CEBS (Wilson et al. 2004; Benek et al. 1996; Breitkreuz and Kennedy 1999; Ziegler 1990). However, thickness maxima (Fig. 3.4.2) along the STZ in northern Denmark (up to 1500 m drilled) as well as along the NNESSW striking Rheinsberg Lineament in NE Germany and NW Poland (up to 2500m of volcanic rocks drilled; Benek et al. 1996) indicate that these zones provided local conduits for the volcanic products. Earliest Permian sediments were deposited in small pull-apart basins between the RFH-TTZ in the north and the EFS in the south (Bachmann and Hoffmann 1997; Gast 1988; Scheck and Bayer 1999) and point to ongoing dextral shearing. Subsequent connection of these initial depocenters led to the establishment of the Northern and Southern Permian Basins (Ziegler 1990) along NW-SE to WNW-ESE-trending axes. These basins were separated by the parallel Mid North Sea-RFZ chain of highs and contain a thick succession of clastic rocks, the Rotliegend, which is overlain by the Late Permian Zechstein evaporites. In the Northern Permian Basin (Fig. 3.4.3), the thick-

Subsidence, inversion and evolution of the thermal field

Figure 3.4.2. Thickness distribution [m] of the Permo-Carboniferous volcanic rocks in the CEBS (modified after Scheck-Wenderoth and Lamarche 2005) showing focused magmatic activity. The hatched line along the axis of the volcanic thickness maximum corresponds to the so-called Rheinsberg Lineament (RL)

Figure 3.4.3. Thickness distribution [m] of Early Permian Rotliegend clastic rocks in the CEBS indicating earlier subsidence in the Southern Permian Basin than in the Northern Permian Basin where the Rotliegend is very thin (modified after Scheck-Wenderoth and Lamarche 2005)

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ness of the Rotliegend rarely exceeds 300 m (Evans et al. 2003; Lockhorst et al. 1998). In contrast, the Southern Permian Basin comprising the Southern North Sea, the North-German Basin, and the Polish Basin, accumulated the thickest Rotliegend sequence in Central Europe with up to 2300 m along the axis of the North German Basin (Bachmann and Hoffmann 1997; McCann 1999; Plein 1995; Scheck and Bayer 1999) and about 1600 m in the Polish Basin (Dadlez et al. 1995, 1998b; Kiersnowski et al. 1995; Lokhorst et al. 1998; ScheckWenderoth and Lamarche 2005). The rather linear structure of the Polish Rotliegend Basin indicates structural control by the crustal-scale Teisseyre-Tornquist Zone in contrast to the adjacent, sag-like basin structure in Northern Germany. Though elongated along the WNW-ESE trend, the latter is wider and two separate thickness maxima indicate two main depocentres. Subsidence analysis across the Southern Permian Basin (van Wees et al. 2000) suggests Late Carboniferous – Early Permian crustal extension followed by delayed infilling and thermal subsidence. The thinner Rotliegend sequence in the Northern Permian Basin documents a later onset of post-rift thermal subsidence than in the Southern Permian Basin. Considerable post-rift subsidence in the Northern Permian Basin is observed in the later Permian Zechstein time (Vejbaek 1997). Seismic and well data suggest that

the Rotliegend is absent on the RFH (Abramovitz and Thybo 2000; Clausen and Pedersen 1999; Scheck et al. 2002b) which represented a structural high during the Early Permian.

3.4.2.2 Late Permian to Early Cretaceous From Late Permian to mid Triassic times, the dominant tectonic regime in the CEBS was cooling of the previously heated lithosphere. Accordingly, post-rift thermal subsidence governed the dynamics of the Polish, North German, and Norwegian-Danish basins though minor normal faulting along predominantly N-S-directed faults indicates some influence of extensional stresses (Clausen and Pedersen 1999; Kiersnowski et al. 1995; Lokhorst et al. 1998; Scheck and Bayer 1999; van Wees et al. 2000; Ziegler 1990). In the Late Permian, a global glacio-eustatic sea-level rise and ongoing rifting in the Greenland-Norwegian Sea established a seaway from the Southern North Sea to the Arctic Ocean (Coward et al. 2003; Ziegler PA 1988, 1990,) opening the way for a marine transgression from the N. This led to the development of the Zechstein Sea with a widening of the depositional area (Fig. 3.4.4) and cyclic sedimentation of evaporites in the central parts of the Northern and Southern Permian basins (Stollhofen et al. this volume; Scheck-

1. Continental basins 2. Deposition of Zechstein salt 3. Faults 4. Suspected basement faults 5. Incipient rifts 6. Variscan Deformation Front

Figure 3.4.4. Palaeogeography of the CEBS during Late Permian Zechstein (~ 255Ma) times. The depositional area widened with respect to the Rotliegend and evaporites were deposited in the Northern and Southern Permian Basins

Subsidence, inversion and evolution of the thermal field

Figure 3.4.5. Cumulative thickness [m] of the Zechstein to Early Cretaceous deposits in the CEBS resulting from the superposition of different tectonic phases: (i) declining Permo-Triassic thermal subsidence causing increasing thickness towards the centres of the NW-SE oriented Northern and Southern Permian Basins, (ii) Mid-Late Triassic extension causing local thickness maxima in N-S-trending grabens and troughs and (iii) Jurassic-Early Cretaceous differential subsidence causing local NW-SE-oriented thickness maxima in the marginal parts of the CEBS. CG: Central Graben, HG: Horn Graben, GG: Glückstadt Graben, LSB: Lower Saxony Basin

Wenderoth et al. this volume). These evaporites played a crucial role during the later Mesozoic structural evolution when the ductile salt layer moved in response to changing stress fields to form a variety of salt structures (ScheckWenderoth et al. this volume). Salt mobilisation caused a complex distribution of salt with different thicknesses. Nevertheless, reconstructions of the initial salt thickness distribution suggest that thermal subsidence was the main controlling factor for basin evolution during the Zechstein (Maystrenko 2005b; Scheck et al. 2003b). The cumulative thickness of Zechstein to Early Cretaceous deposits (Fig. 3.4.5) integrates a long period of differential subsidence and shows two directions of depocentre axes: (i) NW-SE-trending and (ii) N-S-trending. NW-SE oriented thickness maxima are present in the large Norwegian- Danish, Polish, and North German basins (up to 8000 m) as well as in smaller basins along the southern margin of the CEBS in the Broad Fourteens Basin (BFB up to 4000 m), the Lower Saxony Basin (LSB up to 4500 m), and the Subhercynian Basin (SHB up to 3500 m). The largest thickness of the Zechstein to Early Cretaceous layer is present in N-S oriented grabens as in the Central Graben (up to 4500 m), the Horn Graben (up to 6000 m) and the Glückstadt Graben (up to 8500 m).

It has to be considered that this thickness integrates the tectonic events of a large time interval, including phases of salt mobilisation. Furthermore, the presence of subsidence axes sub-perpendicular to each other already indicates that not all of them were active at the same time. In fact, this thickness pattern results from the superposition of at least three different tectonic regimes: (i) declining thermal subsidence causing a regional signal of increasing thickness towards the centres of the Northern and Southern Permian Basins, (ii) Mid-Late Triassic extension causing local thickness maxima in N-S-trending grabens and troughs, and (iii) Jurassic-Early Cretaceous differential subsidence causing local thickness maxima in the marginal parts of the basin system. Though detailed thickness information on the individual subunits composing the above described sequence is not available for the entire basin area, seismic data can help to discriminate the different components. For instance, in the Early Triassic succession of the Danish (Clausen and Korstgard 1993; Clausen and Pedersen 1999; Vejbaek 1997), the North German (Hoffmann and Stiewe 1994; Kockel 1995; Baldschuhn et al. 1996; Kossow and Krawczyk 2002; Scheck et al. 2003a), and the Polish basins (Krzywiec 2002) the declining ther-

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mal subsidence is expressed as continuous and parallel seismic reflections lacking syn-depositional faulting and there is a continuous thickening towards the centres of the Northern and Southern Permian Basins. In contrast, the seismic images are considerably different for the Mid-Late Triassic and indicate syn-depositional normal faulting, local stratigraphic thickening and salt mobilisation in the Glückstadt Graben (Maystrenko et al. 2006), in the Central Graben (Michelsen et al. 1987; Sundsbo and Megson 1993), in the Horn Graben (Best and Kockel 1983; Clausen and Korstgard 1996; Clausen and Pedersen 1999), and in the Rheinsberg Trough (Scheck et al. 2003a). This marked thickening of Triassic deposits in the N-S trending grabens (Fig. 3.4.5) points to a roughly E-W oriented extension (Coward et al. 2003; Scheck-Wenderoth and Lamarche 2005), probably in response to accelerated crustal extension in the Central Atlantic and the Tethys domain (Thierry 2000; Stampfli et al. 2001). This extension also triggered coeval salt rise parallel to the graben margins, thus supporting the creation of accommodation space by salt withdrawal (Scheck-Wenderoth et al. this volume; Maystrenko et al. 2006). Accordingly, this new regime of broadly E-W-oriented extension replaced the thermal subsidence regime which had characterised the evolution of the Permian basins. An exception to this pattern is the MidPolish Trough, where the NW-SE-oriented axis of maximum subsidence remained more or less unchanged because its tectonic evolution was governed by the TTZ (Dadlez et al. 1998b; Krzywiec 2006a). The reduced thickness of the Zechstein to Early Cretaceous on the Mid North Sea-RFH compared to the Danish and North German basins (Clausen and Pedersen 1999; Vejbaek 1997) indicates that it persisted as a structural high during most of the Mesozoic. It was, however affected by progressive dissection as a result of the formation of the north-trending grabens which provided connections between the Norwegian-Danish and North German basins (Scheck-Wenderoth and Lamarche 2005). From Mid-Triassic to Mid-Jurassic times, the southern Eurasian margin was affected by the Cimmerian orogenic cycle which induced far-field intra-plate stresses in the CEBS with associated reduced subsidence and inversion. The Polish Trough, for instance, experienced reduced subsidence or moderate inversion, expressed as depositional gaps or erosional unconformities. On the other hand, the Mid Jurassic was the time when the Central North Sea rift system developed. A deep-seated thermal anomaly beneath the Mid North Sea High is believed to have caused thermal doming in the Central North Sea (Underhill and Partington 1993). By Late Mid-Jurassic times the Central North Sea dome collapsed and subsided again. Though diachronous, Mid-Late Jurassic uplift is regional in scale and has removed much of the initial Early Jurassic succession. In seismic

sections, it is expressed as an erosional unconformity (i.e., Mid-Jurassic or Mid-Cimmerian unconformity) observed from S of the STZ across Denmark to northern Germany, including the Mid-North Sea-RFH and the southern part of the Danish Basin (Vejbaek 1997) as well as the northern half of the North German Basin (Pompeckj Block; Jaritz 1987). While uplift in the Central North Sea and Denmark occurred in the Mid-Jurassic (Surlyk and Ineson 2003; Underhill and Partington 1993), it affected the northern half of the North German Basin (e.g., Pompeckj Block) at the end of the Middle Jurassic (Baldschuhn et al. 1996; Jaritz 1987) and the NE German Basin in the Late Jurassic to Early Cretaceous (Kossow and Krawczyk 2002). Seismic data further show that the situation changed again in the Late Jurassic-Early Cretaceous when narrow, NW-SE-oriented basins subsided along the margins of the CEBS. These are represented by the smaller NW-SE trending thickness maxima in figure 3.4.5. Accordingly, local stratigraphic thickening of the Late Jurassic to Early Cretaceous succession is observed parallel to the STZ at the northern margin of the Norwegian-Danish Basin (Surlyk et al. 2003), in the Polish Basin along the TTZ (MidPolish Trough; Dadlez et al. 1995, 1998b) as well as along the southern margin of the CEBS in the Sole Pit, Broad Fourteens, Lower Saxony and the Subhercynian basins. Coevally, a new phase of salt mobilisation caused salt structures and rim synclines with NW-SW-trending axes to form during the Early Cretaceous (Betz et al. 1987; Jaritz 1980, 1987; Nalpas et al. 1995; Scheck et al. 2002a). This depositional pattern indicates a transtensional regime which again was related to far-field stresses induced by plate-boundary processes such as the northward propagation of the central Atlantic spreading (Stampfli and Borel 2002), extensive rifting in the Norwegian Sea as well as in the Northern and the Central North Sea (Pascal 2002; Torsvik et al. 2002), and early collision phases in the Austro-Carpathian and Balkan areas.

3.4.2.3 Late Cretaceous The Late Cretaceous was a period of elevated sea level, resulting from the high global rates of seafloor spreading (Torsvik et al. 2002). Oceanic spreading in the Central Atlantic propagated northward (Stampfli and Borel 2002; Torsvik et al. 2002) and may have induced minor extensional stresses in the CEBS during early Late Cretaceous times. Together with the high global sea level this resulted initially in flooding of the CEBS during the Late Cretaceous and in deposition of carbonates seamed with clastic rocks along the southern margin. During the latest Cretaceous, orogenic activity along the southern Eurasian margin moved from the Austro-Carpathian in the E to the Alpine Tethys area and extended westwards into the

Subsidence, inversion and evolution of the thermal field

Figure 3.4.6. The present-day thickness distribution of Late Cretaceous in the CEBS reveals a dominance of NW-SE striking, inversionrelated structural elements. Uplifted highs at the northern and southern margins of the CEBS as well as in the Polish Basin are bordered by parallel troughs. The present distribution of Late Cretaceous is the combined geometrical result of inversion, erosion and sedimentation

western Alps in the Late Cretaceous (Stampfli and Borel 2002). This convergence induced compressive stresses in the lithosphere of the CEBS, and consequently, early Late Cretaceous subsidence was followed by a period of inversion, when in particular the marginal parts of the Norwegian-Danish and the North German Basins as well as the axial part of the Polish Basin were affected by uplift. In response to this regime, the thickness distribution of the preserved Late Cretaceous chalk in the CEBS indicates a clear dominance of NW-SE structural elements (Fig. 3.4.6) and reflects the localised uplift along the STZ, along the TTZ and along the EFS at the southern margin of the CEBS. Most of the NW-SE-striking inverted blocks in the CEBS are bordered by marginal basins which are truncated by the Base Cenozoic Unconformity - an indication that uplift occurred prior to Cenozoic deposition. The areas affected by Late Cretaceous-Early Cenozoic uplift coincide with depocentres in the preceding Late Jurassic-Early Cretaceous period. Many of the reverse or transpressive faults accommodating uplift show normal or transtensional activity during the deposition of the underlying Late Jurassic-Early Cretaceous sediments, as for example in the Lower Saxony, Sole Pit (Badley et al. 1993; Buchanan et al. 1996; Nalpas et al. 1995), Broad

Fourteens (De Lugt et al. 2003; Nalpas et al. 1995), and the Subhercynian basins (Franzke et al. 2004; Kossow et al. 2001; Otto 2003; Voigt et al. 2004). In the North Sea, inversion was most intense along the NW-SE-striking branch of the Central Graben and appears to be synchronous (late Hauterivian into Palaeogene) with inversion along the Sorgenfrei-Tornquist Zone at the northern margin of the Norwegian-Danish Basin (Cartwright 1989; Vejbaek and Andersen 2002). In both areas, the location of inversion is spatially linked to Late Jurassic-Early Cretaceous, NW-SE-trending depocentres (Vejbaek 1997). Likewise, Late Cretaceous inversion of the Mid Polish Basin took place along the axial part which was the area of maximum subsidence in Permian to Early Cretaceous times (Fig. 3.4.6; Erlström et al. 1997; Krzywiec 2002; Krzywiec et al. 2003). Though deep erosion on this Mid Polish Swell renders it difficult to determine the timing of the inversion, evidence from the basin margins suggests a time window between the late Turonian and the earliest Tertiary (Dadlez 1994; Krzywiec 2002; Krzywiec et al. 2003). Inversion was accompanied by a new period of salt movement during which salt diapirs with NW-SE trending structural axes formed parallel to the uplifted blocks

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and again decoupled deformation in the sub-salt basement from the salt cover (Scheck-Wenderoth et al. this volume). The orientation and geometry of inversion structures (NW-SE) as well as palaeo stress studies (Nalpas et al. 1995; Franzke et al. 2004; Kossow et al. 2001; Voigt et al. 2004; Mazur and Scheck-Wenderoth 2005; Vandycke and Bergerat 2001; Lamarche et al. 2002; Lamarche et al. 1999; Sippel at al. in press) suggest that the CEBS underwent N-S to NE-SW compression during the Late Cretaceous. This is parallel to the axes of older N-S-striking depocentres and could explain why the latter have generally not experienced inversion. Instead, the N-S grabens are covered with a more or less constant thickness of Late Cretaceous (Vejbaek and Andersen 2002; Maystrenko et al. 2005a). Thus, the CEBS was first flooded in the early Late Cretaceous and later subjected to compressive deformation localized predominantly along NWSE-striking fault zones of inherited crustal weakness. The Cretaceous-Tertiary transition is expressed as an erosional unconformity over the largest part of the CEBS and represents a renewed change in its tectonic history.

3.4.2.4 Cenozoic The thickness distribution of the Cenozoic (Fig. 3.4.7) indicates localised subsidence in northerly trending troughs in contrast to the dominant NW-SE-direction in the Late

Cretaceous, and the major depo-centre with up to 3500 m of Cenozoic sediments is shifted to the Central North Sea. A small NW-SE striking maximum is present in the Roer Valley in The Netherlands. The outlines of a larger, NWSE oriented depo-centre in Northern Germany, The Netherlands, and Southern Denmark are visible as well, with an average Cenozoic thickness of about 1000 m. Within this depression, local thickness maxima of up to 2100 m again strike N-S. Several N-S trending structures that had been tectonically quiet during the phase of Late Cretaceous inversion experienced extensional reactivation in the Cenozoic. A renewed phase of salt rise parallel to the margins of these northerly trending depocentres and coeval normal faulting additionally indicate that the stress field controlling subsidence had changed with respect to the Late Cretaceous.

3.4.2.5 Summary Within the CEBS, the STZ, the RFH and the EFS bordered the large Permo-Mesozoic Norwegian-Danish and North German Basins which experienced Permian-Mesozoic subsidence along axes parallel to these fault systems after thermal destabilisation of the crust in the Late Carboniferous-Early Permian. The most intense subsidence in the Polish Basin occurred along the TTZ. A change of the regional stress field in the Triassic caused subsidence of

Figure 3.4.7 Thickness distribution [m] of Cenozoic deposits in the CEBS (modified after Scheck-Wenderoth and Lamarche 2005) indicating a shift in deposition towards the North Sea and a general dominance of N-S trending depocentres

Subsidence, inversion and evolution of the thermal field

Table. 3.4.1. Comparison of compilations of thermal conductivities for important lithologies ”in W/m/K“ PetroMod Lithology Shale Shale coaly Siltstone Sandstone Conglomerate Coal Limestone Marl Dolomite Salt Anhydrite Water Oil Methane

20°C

100°C

1.98 1.8 2.14 3.12 2.93 0.5 2.83 2.23 3.81 5.69 4.81

1.91 1.6 2.03 2.64 2.63 0.46 2.56 2.11 3.21 4.76 3.97

Beardsmore* (1996)

Beach et al. (1987)

Reiter and Tovar (1982)

Barker (1996)

2.9

1.4 ± 0.4

2.1 ± 0.4

1.8 ± 1.2

2.9 7.1

3.2 ± 1.3 3.1 ± 1.3 3.2 ± 1.8 0.2 ± 0.2 2.4 ± 0.9 3.0 ± 1.1 3.1 ± 1.4 5.7 ± 1.0

2.7 ± 0.2 3.7 ± 1.2

4.7 ± 2.8

3.1 3.2

2.8 ± 0.3 4.7 ± 1.1 5.4 ± 0.3

Poelchau et al. (1997)

2.5 ± 0.6 2.4 ± 0.5 3.7 ± 1.8 5.9 0.6 0.15 0.03

* Matrix conductivities only representing rock conductivities when porosity is zero.

roughly N-S striking grabens (e.g., Central Graben, Horn Graben, Glückstadt Graben) roughly perpendicular to the STZ, RFH and EFS. The resulting subsidence pattern can be explained by a superposition of declining thermal subsidence and regional, E-W directed extension. Far-field stresses caused further differentiation of the CEBS during the Jurassic to Early Cretaceous into local sub-basins and structural highs. In the Late Cretaceous-Early Cenozoic, far-field compression resulted in inversion of the Norwegian-Danish and North German basins with localised uplift along NW-SE-oriented marginal fault zones (STZ, RFH and EFS) and along the TTZ in the Polish Basin, whereas N-S- trending structures experienced almost no inversion. A renewed change in subsidence controlling factors led to a shift of the main depocentre to the central North Sea and to extensional reactivation of the N-Strending Triassic graben structures in the Cenozoic.

3.4.3 Temperature in sedimentary basins The temperature field of sedimentary basins is one of the decisive factors governing oil and gas generation, the conversion of peat into lignite and coal as well as many other diagenetic reactions. Also several petrophysical properties depend to a great extent on temperature. Finally, geothermal energy can be produced from sedimentary rocks under ideal conditions. Therefore, understanding of the present and past temperature field is of great importance in earth sciences. For details on the theory of heat transfer in sedimentary basins, the publications by Yalcin et al. (1997) and Beardsmore and Cull (2001) are recommended.

Temperature distribution depends basically on three processes, (i) heat conduction, (ii) heat convection, and (iii) radiation. Electromagnetic radiation of the sun greatly influences annual mean surface temperature of solid earth and has some influence on the internal temperature of the upper part of sedimentary basins. Global mean surface palaeotemperatures have been published by Frakes (1979) and mean surface palaeotemperatures for differrent latitudes have been compiled by Wygrala (1989). In the deep sea, surface temperatures also depend on water depth. They are less than 5 °C at water depths of more than 1500 m, independent of sea surface temperatures. Conductive heat transport is a diffusive process that takes place in all sedimentary basins where it is commonly the dominant process for internal heat transfer. The rate of vertical heat flow is often expressed as Q= -λ · δT/δz

(3.4.1)

in which Q is the heat flow (mW/m2) and λ is the thermal conductivity (W/m/K). Positive heat flow is conventionally taken for the direction of decreasing temperature, whereas the convention of geothermal gradient is that it is positive in the direction of increasing temperature. This is the reason for the negative sign in the equation 3.4.1 which is often neglected (see Beardsmore and Cull 2001 for more details). The equation implies that the temperature gradient in a sedimentary basin is correlated to heat flow –grad T=Q/λ

(3.4.2)

but that it also depends on thermal conductivity. Thermal conductivity, however, is (i) highly variable in sedi-

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mentary systems, (ii) porosity-dependent, and (iii) temperature-dependent. Therefore, there is no “uniform” or “linear” geothermal gradient in the Earth’s crust, even if it is composed of a homogeneous lithology and even if convective processes are absent. Thermal conductivities of sedimentary rocks are listed in table 3.4.1 based on data summarised in Yalcin et al. (1997), Beardsmore and Cull (2001), and the user manual of the commonly applied basin modelling software PetroMod. Thermal conductivity decreases with increasing porosity because pores are usually filled with water, oil, or gas, which all have a low thermal conductivity (Table 3.4.1). Porosity loss due to compaction would therefore lead to increasing thermal conductivities of rocks with depth. On the other hand, thermal conductivity is temperature-dependent and decreases with increasing temperature. This effect leads to decreasing thermal conductivities with depth. The bulk effect of these counteracting sensitivities has been summarised by Ungerer et al. (1990) and is visualised in figure 3.4.8. In many cases, mixed lithologies exist, for which no adequate thermal conductivity measurements (at different temperatures) are available. In this case the average conductivity has to be calculated based on the relative proportions of pure lithologies. Examples from the CEBS are summarised in table 3.4.2. The principles and pitfalls

of such calculation have been discussed in detail by Hartmann et al. (2005). Convective heat transport is basically a mass transport achieved by fluid flow. Fluids can be water, which is the most common case, oil or gas. The moving fluids can be cool or hot and thus have a cooling or heating effect on sedimentary basins. Examples of hot fluids entering sedimentary systems include hydrothermal water related to magmatic activity and also fluids moving upwards in sedimentary basins from great depth into shallower levels (Bethke 1989; Wycherley et al. 2003). In contrast, downward movement of meteoric waters from the surface into deep basin parts has a pronounced cooling effect. Such forced convection is most effective in basins where a steep surface morphology exists and/or where open faults perturb the thermal convection flow pattern. Examples include both foreland and rift basins (Simms and Garven 2004). In contrast, compaction-driven expulsion of formation water that is not focussed along highly permeable layers is generally too slow to affect the temperature field significantly (Hermanrud 1986). Heat is generated not only in the lower crust and mantle, entering sedimentary basins from the base (basal heat flow), but also inside the basins. This depends on the concentration of radioactive elements which is highly variFigure 3.4.8. Vertical thermal conductivities of porous rocks during burial, as calculated for two different thermal gradients (modified from Ungerer et al. 1990)

Subsidence, inversion and evolution of the thermal field

Table. 3.4.2. Thermal conductivities of several stratigraphic units in the east-German part of the CEBS ”in W/m/K“ (Lotz 2004) Shale Stratigraphy Keuper Muschelkalk Buntsandstein Zechstein Elbe Subgroup Havel Subgroup Müritz Subgroup Altmark Subgroup Stefanian Westphalian Namurian Visean Frasnean Givetian

Siltstone

λ⊥

λP

λ⊥

λP

Sandstone λ⊥

Conglomerate

λP

λ⊥

λP

3.5 3.3 4.0 4.2 2.6

4.1 4.4 2.8

3.2 3.3 3.2 3.4

3.1 3.2 3.3

4.7 4.6 3.6

5.1 4.8 3.9

Lime - / Marlstone λ⊥ λP

Anhydrite / Dolomite λ⊥

λP

3.0

3.0

5.3

5.3

2.5 2.5

2.6 2.8

2.9 2.3 2.3 0.9 2.3 2.8

3.6

2.3 2.5 1.5 2.7 3.1

3.8

2.9 2.9

3.2 3.6

2.3

2.7

3.4 3.0

3.7

2.7 4.8

Eifelian

3.6

4.7

3.9

Table. 3.4.3. Average contents of heat producing radioactive elements in sedimentary rocks (after Rybach 1986) Rock Type

U [ppm]

Th [ppm]

K [%]

Th/U

Limestone

2.0

1.5

0.3

0.75

0.62

Dolomite

1.0

0.8

0.7

0.80

0.36

Salt

0.02

0.01

0.1

0.50

2.2

0.01

Anhydrite

0.1

0.3

0.4

3.0

2.9

0.09

Shales, siltstones

3.7

12.0

2.7

3.2

2.4

1.8

Black shales

20.2

10.9

2.6

0.54

2.4

5.5

Quartzite

0.6

1.8

0.9

3.0

0.32

Arkose

1.5

5.0

2.3

3.3

0.84

Graywacke

2.0

7.0

1.3

3.5

Deep sea sediments

2.1

11.0

2.5

5.2

Carbonates

Density* [10³ kg/m³]

Heat generation [µW/m³]

2.6

Evaporites

Sandstones

2.4

0.99 1.3

0.74

* Broad average since density strongly depends on porosity.

able in different rocks. Rybach (1986) calculated the effect of heat production on the thermal field by a simple, purely conductive, one-dimensional model. Temperature is calculated for any depth (hz) as: T(h) = T0 + hz(Q + AH)/k – h2zA/(2k)

at the base, A is the average radiogenic heat production, k is the average thermal conductivity, and H is the thickness (see Yalcin et al. 1997). As heat generated by internal sources is also dissipated by conduction, the temperature increase is also controlled by thermal conductivity.

(3.4.3)

where T0 is the surface temperature, Q is the heat flow

A summary for different sedimentary rocks has been compiled in table 3.4.3 based on Rybach (1986). In gen-

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Figure 3.4.9. Effect of heat generation (A) in sediment and thermal conductivity (k) on a temperature-depth profile. Temperature curves are calculated with the given k and A values and from equation 3.4.3 for H = 6 km and with Q = 70 mW/m2 and T0 = 10 °C (after Rybach 1986)

eral, heat generation is high in coals and oil shales containing incompatible elements, at medium level in shales and siltstones containing abundant potassium, and low in carbonates, sandstones and evaporites. For most sedimentary basins, internal heat production is at only 10% of the basal heat flow (and often much less), but there are exceptions. Such exceptions can be found in clay-rich basins, and also in basins intruded by granitic magmatites. The effect of heat production is visualized in figure 3.4.9 (after Rybach 1986).

Such rapid geological processes have to be taken into account in any temperature history reconstruction.

Old sedimentary basins are usually in a state of thermal equilibrium, if they are not affected by active tectonics. This is, however, not the case in areas with rapid temperature changes over time. Such a situation exists, for example, along subduction zones which are common on earth, having a total length of 43,500 km at present day. There, cool sediments are rapidly transported to great depth affecting the temperature field. The example of the continental margin of Costa Rica is shown in figure 3.4.10 (after Lutz et al. 2004).

The present-day temperature field in the CEBS has been studied in great detail over the last decades. Hurter and Haenel (2002) published maps of subsurface temperatures and heat flows. Based on their data, a simplified map has been compiled for the CEBS (Fig. 3.4.11). Heat flows are at a moderate level for continental crust (Allen and Allen 2005), but slightly elevated at e.g., the area north of Berlin. Along the southwestern border, there is a zone of warm springs, where heat flows are higher than average.

Present-day heat flows cannot be directly measured but are calculated from temperatures measured in the subsurface and thermal conductivities. Heat flows vary significantly and are commonly high in areas of volcanic activity and in areas of active extensional tectonics (Sclater et al. 1980). Examples of areas of high heat flow in Europe are Iceland and the Upper Rhine Valley.

Subsidence, inversion and evolution of the thermal field Figure 3.4.10. Calculated present-day temperature distribution for two different subduction velocities (A, 120 mm/a; B, 90 mm/a) calculated for the continental margin of Costa Rica. Colours indicate isotherms. GP: Grid points in the model (modified from Lutz et al. 2004)

3.4.4 Maturity and temperature parameters in sedimentary basins Whereas the present-day temperature field inside sedimentary basins can be studied on the basis of borehole temperatures or temperature logs, the reconstruction of palaeotemperature evolution is more difficult. In many sedimentary basins, present temperatures are much lower than ancient temperatures and the state of the rocks with respect to, e.g., petroleum generation and compaction was established during earlier geologic eras at higher temperatures. Especially for these basins, understanding of the palaeotemperature histories is critical and a prerequisite when quantifying diagenesis or petroleum generation. However, palaeotemperature reconstructions are essential not only for basins and sedimentary rocks which experienced their highest temperatures in the past, but also for those in which Neogene temperatures are/were the highest. This is because early temperature evolution will have already influenced mineral precipitation and oil or gas generation from petroleum source rocks. The knowledge

of the extent of such an early phase of generation can be a clue towards an understanding of the extent of the late (Neogene) phase of petroleum generation in different parts of a basin and thus an aid in exploration strategies. The best way to reconstruct temperature histories in the context of basin history is numerical modelling (see below). The quality of model predictions (on palaeotemperature, petroleum generation etc.) depends on the availabilty of temperature-sensitive data and parameters, which can be used for calibrating the model. These parameters are roughly subdivided into organic maturity parameters and inorganic temperature parameters. Maturation is a term commonly used in sedimentary basin studies to address thermally induced changes in the nature of organic matter. Maturation depends on temperature and the time, during which specific temperatures are maintained. Other factors such as the chemical environment and pressure are generally regarded to be of lesser importance, although there may be exceptions (Carr 1999; Huang 1996; Price and Barker 1985). In order to quantify

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Figure 3.4.11. Map of present-day heat flows in the CEBS (based on Hurter and Haenel 2002 and Królikowski 2006)

thermal maturation of sedimentary rocks, a great number of physical and chemical maturity parameters has been developed. All these parameters are measured either on total organic matter or parts of the organic matter. Table 3.4.4 gives an overview on some of the most common parameters used and their application. The most widely used parameters in the petroleum industry are vitrinite reflectance, spore colour, and Tmax-values. Furthermore, molecular geochemical parameters are applied, which are especially useful in oil-oil and oil-source rock correlations.

burial and temperature models described below are generally based on calibration using the Sweeney and Burnham (1989) algorithm. However, some researchers argue that under “normal” burial conditions there is always sufficient time available for vitrinite particles to adapt and that, therefore, vitrinite reflectance can be directly translated into maximum palaeotemperature (Barker and Pawlewicz 1994). They developed

Vitrinites are organic particles derived from higher plants (“wood-like particles”) which are the major constituents of most coals, but also ubiquitous in other sedimentary rocks. Their chemical properties as well as their reflectance change systematically as a function of temperature and time. Vitrinite reflectance has accordingly been widely applied in order to quantify thermal histories (Taylor et al. 1998). The evolution of vitrinite reflectance as a function of temperature and time has been intensely studied for more than 50 years. At present, calculation of temperature histories from vitrinite reflectance data is mainly based on the algorithm published by Sweeney and Burnham (1989):

for “normal” burial conditions (slow heating) and for hydrothermal conditions (rapid heating):

VRr= e (-1.6 + 3.7F)

(3.4.4)

In this equation, F is a stochiometric factor ranging from 0 to 0.85 and VRr is the mean vitrinite reflectance measured in oil immersion on randomly orientated grains. The

Tpeak = (lnVRr + 1.68) / 0.0124

(3.4.5)

(3.4.6) Tpeak = (lnVRr + 1.19) / 0.00782 Vitrinite reflectance ranges from 0.2 to 0.5% at the immature stage which is roughly equivalent to the peat, lignite, and subbituminous coal stage, from 0.5 to 1.3% in the mature petroleum generation stage, from 1.3 to 2.0% in the wet gas stage and is above 2% in the dry gas stage which is roughly equivalent to anthracite rank. Aside from vitrinite reflectance, many other optical maturity parameters have been developed (Table 3.4.4). Solid bitumen reflectance proved to be particularly useful, since this type of particle is common in many rocks in petroleum-bearing basins. Solid bitumen is derived from decomposition (cracking) of former oil and is accordingly most common in petroleum source and reservoir rocks, but also in former carrier rocks. These rocks are often poor in vitrinite; therefore

Subsidence, inversion and evolution of the thermal field

Table 3.4.4. Comparison of different maturity parameters Maturation Parameter

Method

Short Term(s)

Application

Range

Vitrinite Reflectance

Optical, selected woody organic particles

VRr; Rr; Ro

Devonian and younger

immature (0.2-0.5%), mature (0.5-1.3%), slightly overmature (1.3-2.0%) and highly overmature (>2.0%)

Maximum Vitrinite Reflectance

Optical, selected woody organic particles

Rmax; VRmax

Devonian and younger

Semi-Anthracite (2.2%) to Metamorphism >6%)

Solid Bitumen Reflectance

Optical, selected organic particles

BRr

All sed. rocks, preferentially source,- carrier-, and reservoir rocks

mature to highly overmature

Fluorescence Colour and Intensity

Optical, selected organic particles, e.g., spores, algae

λmax, Q, Flint

All sed. rocks, preferentially source rocks; also oil inclusions

immature to mature

Spore Colour

Optical, spores, pollen grains

TAI

Devonian and younger, preferentially terrigenous environments

immature to mature

Conodont Colour

Optical, conodonts

CAI

Marine Rocks

immature to mature

Carbon Preference Index

Geochemical, specific non-aromatic hc

CPI

All sed. rocks, unweathered material required

immature to mature

Methyl Phenantrene Index

Geochemical, specific aromatic hc

MPI

All sed. rocks, unweathered material required

early mature to slightly overmature

“Biomarkers” (Vari-

Geochemical, specific non-aromatic or aromatic hc

See Peters et al. (2005)

All sed. rocks, unweathered material required

immature to mature

ous Parameters) Rock-Eval Tmax

Geochemical, total organic matter

Tmax

All sed. rocks

immature to slightly overmature

Rock-Eval PI

Geochemical, total organic matter

PI

All sed. rocks, unweathered material required

immature to slightly overmature

Rock-Eval HI

Geochemical, total organic matter

HI

All sed. rocks

Volatile Matter Yield

Geochemical, total organic matter (water- and ash-free)

VM

Only Coals

immature to slightly overmature, depends on petroleum generation potential and maturity immature to overmature

Water content/moisture

Geochemical, total organic matter (ashfree)

-

Only Coals

immature to early mature

Carbon content

Geochemical, total organic matter (water- and ash-free)

C

Only Coals

immature to overmature

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reflectance of solid bitumen can act as an alternative maturity parameter there. It also occurs in rocks older than Devonian in which vitrinite particles sensu strictu are absent due to the lack of higher land plants. Correlation between vitrinite and solid bitumen reflectance is visualised in figure 3.4.12a (Schoenherr et al. 2007a) and can be expressed as

VRr =

BRr + 0.2443 1.0495

(3.4.7)

Further optical maturity parameters are listed in table 3.4.4. In particular, spore colour and conodont colour as well as fluorescence parameters can be regarded as rapid indicators for thermal maturation, but are not very reliable as quantitative calibration tools in thermal modelling because colour/fluorescence depend also on other factors rather than just temperature, e.g., on the thickness of the fossil or on the degree of oil impregnation. Excellent maturity parameters exist for coals, including vitrinite reflectance, water content, volatile matter yield, and carbon content (Table 3.4.4; Taylor et al. 1998). Aside from these parameters, calorific values are also used which are of direct relevance for coal utilisation in power plants. With the exception of vitrinite reflectance, none of the above parameters can be utilised on rocks other than coal. In coals, maturity is often called “rank” or “coalification”.

a

In contrast, there are several geochemical parameters applicable for different rock types. In particular, RockEval parameters Tmax (temperature of maximum pyrolysis yield) and PI (Production Index) have often been used in the petroleum industry. PI values depend on the progress of petroleum generation and tend to increase with depth/temperature, but are also influenced by petroleum impregnation or expulsion. The more widely used Tmax values increase with depth/temperature. They have been correlated with vitrinite reflectance for terrestrial organic matter (type III kerogen) and humic coals (Teichmüller and Durand 1983; Fig. 3.4.12b), but depend to a great extent on thermal stability of organic matter. Nevertheless, Tmax values in combination with Hydrogen Index (HI) and PI values from Rock-Eval pyrolysis have excellent potential to predict quality of organic matter (kerogen) with respect to petroleum generation potential and maturity. Pitfalls of the technique have been discussed in detail, e.g., by Dahl et al. (2004) and Peters (1986). In particular, the presence or absence of specific minerals can change Rock-Eval parameters. One common effect is that hydrocarbons released from kerogen are hindered from migrating out of the source rock, leading to lower HI and higher Tmax values. Even greater effects exist for certain ironbearing minerals which react with hydrocarbons, leading to greatly changed Rock-Eval parameters. Examples from a series of experiments are shown in figure 3.4.13. Comparison of Rock-Eval parameters with a variety of other

b

c

Figure 3.4.12. Correlation of vitrinite reflectance with solid bitumen reflectance for two large data sets (a); Schoenherr et al. 2007a Tmax values from Rock-Eval pyrolysis (b) Taylor et al. 1998 and MPI (c) Radke et al. 1997; G.R.S. is Green River Shale; see Table 3.4.4

Subsidence, inversion and evolution of the thermal field

coal rank parameters listed in table 3.4.4 has been published by Bostick and Daws (1994). In addition, there are many molecular geochemical parameters which are well suited to study thermal maturity, especially for the range of maximum temperatures between 50 and 150 °C. These parameters are usually less suitable for weathered rocks from outcrops and for high ranges of maturity, exceeding 1.3% VRr. An exception is the Methylphenantrene Index (MPI) established for terrigenous type III kerogen by Radke and Welte (1983), which covers a greater maturity range than other molecular parameters (Fig. 3.4.12c). Also widely applied is the “Carbon Preference Index” (CPI) value which depends on the ratio of odd-numbered over even-numbered n-alkanes. In addition, there are many parameters which depend on ratios of similar molecules which have, however, different thermal stabilities. Details on geochemical maturity parameters can be found in Peters et al. (2005). Inorganic temperature parameters commonly applied in sedimentary systems include fission track data, fluid inclusion data, and – less commonly - clay mineral transformation ratios (Velde 1995). Fluid inclusions are treated in Gaupp et al. (this volume) and Schöner et al. (this volume). The method of fission track analysis is described in box 3.4.1.

3.4.5 Variability of palaeotemperature fields in the Central European Basin System This section describes the regional thermal evolution of the basin, considering some cross-sections in their actual and former configurations (Fig. 3.4.14). Section A-A’ is situated at the south-western margin of the CEBS in the Netherlands. A detailed description of burial and temperature histories for individual wells is found in Nelskamp et al. (2007). There, as well as in other parts of the basin, little is known about Early Carboniferous and older deposits. The Late Carboniferous is represented by coal-bearing strata, acting as the most important gas source rocks in the basin. During the Permian, sedimentation was controlled by the edges of the Southern Permian Basin (SPB). Rotliegend sediments were only deposited in the Central Netherlands Basin (CNB), while Zechstein sediments were deposited along the whole section with the exception of the London-Brabant Massif (LBM) in the southwestern part. The thickness of Zechstein deposits decreases towards the LBM (Fig. 3.4.15) and salt deposits can only be found in the north-eastern most part of the section. Further to the north and towards the basin centre, Zechstein deposits of more than 900 m can be found (Duin et al. 2006). The thickness of Early Triassic sediments increases from the

Figure 3.4.13. Rock-Eval Tmax and HI parameters of a lignite sample mixed with different pure minerals

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Box 3.4.1 Fission track thermochronology Figure1. Fission tracks in zircon. Different shapes and cross sections as seen in the focal plane result from different inclinations with respect to the polished surface. A horizontal confined track, not intersecting the polished surface of the crystal, is indicated by arrows and displayed for two different focal planes. The long edge of the photo measures approximately 30 microns. Fission track (FT) thermochronology provides information on the thermal history of a rock. In sedimentary basins, this usually in the first place means amount and timing of heating during subsidence and in the second place cooling related to uplift and/or denudation after inversion (e.g., Senglaub et al. 2005). FT data, however, may also help to characterize and differentiate provenance areas by specific crystal age spectra (e.g., Bernet and Spiegel 2004). Dating the formation ages of volcanic layers (e.g., Gleadow 1980) may contribute to the set up of an absolute time frame within a basin. Finally, FT analysis may define constraints for the cooling of plutonic or metamorphic basement and thus for the start of basin development. FT analysis has been developed as a radiometric method based on the spontaneous fission of 238U (e.g., Wagner and van den Haute 1992). Moving in opposite directions through the crystal lattice of the host mineral, the two fission fragments capture electrons from the atoms passed by and cause linear zones of lattice damage. These zones, the fission tracks, may be stable over geological times and may thus be regarded as daughter effects comparable to the daughter isotopes of classical radiometric dating methods. The tracks are usually studied in apatite, zircon (Fig. 1), and titanite. They may be revealed by etching so that they become visible under an optical microscope. The calculation of an age, under fast cooling conditions representing the time elapsed since the last cooling below a temperature specific for each mineral, is primarily based on the ratio given by the U content and the number of tracks accumulated within a certain volume of a crystal. Figure 2. Schematical illustration of the length distributions of horizontal confined tracks, resulting from thermal histories, passing in specific ways through the partial annealing zone of fission tracks. Beyond the scope of classic radiometric dating methods, providing purely numerical age information, FT data reflect as well the temperature path experienced by the studied rock. This results from the fact that there is no sharp closure temperature for fission tracks, but instead a temperature range of partial track stability. Within this range (called PAZ, partial annealing zone), the length of the tracks is successively being shortened. The resulting track length distributions are representative for specific temperature histories (Fig. 2), e.g., slow cooling or (re-)heating. Within the PAZ, also the number of tracks may be reduced. The resulting apparent age does no longer represent the primary depositional age of the sample. The importance of FT analysis for hydrocarbon exploration results from the fact that the PAZ of FT in apatite roughly coincides with the oil window during hydrocarbon maturation and therefore provides valuable informations for decisions on exploration strategies. Hypothetical evolutions of track lengths and ages may be calculated using genetic algorithms and geologically meaningful boundary conditions. Relevant FT parameters (i.e., apparent age, mean, standard deviation, and skewness of the length distribution of horizontal confined tracks) resulting from different temperature paths and boundary conditions may be compared to the observed values. Those with the best fit are generally assumed to reflect the history of the studied rock (Fig. 3, thick line). Those with a reasonable fit may point at alternative interpretations. The most commonly used programs are described in Gallagher (1995), Dunkl (2002), and Ketcham et al. (2003). Figure 3. Modelled thermal evolution of a sample, based on the comparison of calculated evolutions of FT parameters, using genetic algorithms, with measured data; thick line and red area = good fit = probable evolution; green area = reasonable fit = alternative interpretations; white area = bad fit = unrealistic assumptions; for details see text.

Subsidence, inversion and evolution of the thermal field

Figure 3.4.14. Location map of seismic profiles for which numerical simulations on burial, erosion and temperature histories are presented. The blue area represents the distribution of the Zechstein. Structural elements: CNB - Central Netherlands Basin, EEC - East European Craton, EFZ - Elbe Fault Zone, FP - Friesland Platform, LBM - London Brabant Massif, RVG - Roer Valley Graben, WNB - West Netherlands Basin, ZR - Zandvoord Ridge

northeast to the southwest. In the Late Triassic, erosion occurred on several structural highs, e.g., on the Friesland Platform, where locally the entire Triassic succession was removed. The West Netherlands Basin (WNB) was formed further to the southwest (Fig. 3.4.15, lower part). During the Early and Middle Jurassic, the whole area was flooded and thick successions of sediments were deposited in the WNB and CNB. In the Late Jurassic/Early Cretaceous, tectonic movements caused erosion which removed the entire Jurassic section on the highs and most of the Middle Jurassic section in the basins (Nelskamp et al. 2007). After this erosion phase, sedimentation started again in the basins. A Cretaceous transgression flooded the whole area and chalk was deposited. In the Late Cretaceous and Early Tertiary, the so-called Subhercynian/Laramide Tectonic Phase caused uplift and inversion of the basins. The WNB was tilted, erosion being most pronounced in the northwest along the Zandvoort Ridge. Up to 2000 m of Triassic to Cretaceous sediments were removed, contrasting to only 200 to 500 m in the southwest (Fig. 3.4.15, middle part). The CNB was inverted along its central axis. Erosion was therefore greatest in the centre of the basin where locally all sediments from the Carboniferous to

the Cretaceous were removed, in total up to 2500 m. In the Netherlands section of the LSB, the inversion movements were only small and removed only up to 500 m. Thick Late Cretaceous sediments are preserved on some of the former highs (up to 1800 m, Fig. 3.4.15, upper part). In the Tertiary, renewed sedimentation was again interrupted by another inversion event at the end of the Eocene, which is related to the Pyrenean tectonic pulse (de Jager 2003). The WNB was more strongly affected by this erosion than the CNB. All sediments from the Palaeocene and Eocene were eroded in some places. Further to the east, the Lower Saxony Basin (LSB) developed in the Late Jurassic and Early Cretaceous (Figs. 3.4.14: section B-B’, 3.4.16) with thick deposits. Thermal subsidence ceased in the course of the Early Cretaceous and was followed by inversion and erosion during the Late Cretaceous (Fig. 3.4.16, middle part). Strongest inversion/ erosion occurred in the former basin centre, whereas it was much less pronounced towards the north. In total, up to 6 kilometres of sediments, locally even more, were eroded during this phase (Senglaub et al. 2006; Adriasola-Muñoz et al. 2007; Voigt et al. this volume). As a consequence,

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R. Littke · M. Scheck-Wenderoth · M.R. Brix · S. Nelskamp Figure 3.4.15a. Stratigraphy of profile A-A’ through the western part of the basin (Netherlands). Three time steps are shown: Late Jurassic, MidCretaceous (inversion phase) and present-day. The legend is also valid for figures 3.4.15b, 3.4.16, 3.4.17 und 3.4.18

Subsidence, inversion and evolution of the thermal field Figure 3.4.15b. Temperature evolution of profile A-A’ through the western part of the basin (Netherlands). Three time steps are shown: Late Jurassic, Mid-Cretaceous (inversion phase) and present-day

highly compacted and thermally highly mature sedimentary rocks are present at the Earth’s surface now. The transition between the LSB and the Pompeckj Block further north is shown in section C-C’ (Fig. 3.4.14; 3.4.17; Schwarzer and Littke 2007). Compared to the areas further to the west discussed previously, much thicker Zechstein deposits top the coal-bearing Late Carboniferous and the thin Rotliegend strata, which include important gas reservoir sandstones. Here and further north in the Schleswig-Holstein area (Maystrenko et al. 2005a), salt tectonics greatly influenced the sedimentation pattern. A quite uniform sedimentation prevailed from the Permian to the Middle Jurassic, although Late Triassic deposits were thicker in the (southern) LSB part. Subsidence changed dramatically in the Late Jurassic, when

uplift and erosion affected the Pompeckj Block, whereas strong subsidence and erosion occurred in the southern part (LSB). This pattern continued during the Early Cretaceous, when the rocks of the LSB reached their greatest depth. Thereafter, Late Cretaceous inversion removed hundreds to thousands of metres of sediments in the LSB (Voigt et al. this volume), and strong subsidence and sedimentation occurred in the northern Pompeckj Block. As a consequence, the Zechstein base is now surprisingly flat (Fig. 3.4.17), but it wasn’t in the past! Further to the east, the coal-bearing Late Carboniferous sequence is missing in the south, where Namurian and older rocks are overlain by a very thick sequence of Rotliegend rocks (Section D-D’ in Fig. 3.4.14; Fig. 3.4.18, lower part). Rotliegend rocks are divided into a lower se-

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R. Littke · M. Scheck-Wenderoth · M.R. Brix · S. Nelskamp Figure 3.4.16. Profile B-B’ through the Western Lower Saxony Basin (Bramsche anomaly). Three time steps are shown: Early Cretaceous (Wealden), Mid-Cretaceous (inversion phase) and present-day

quence of predominantly volcanic rocks which witness the initial phase of the CEBS (Breitkreuz et al. this volume) and an overlying sedimentary sequence. The evaporitic Zechstein formation is thick in that area, although salt diapirism is less pronounced than further towards the northwest. Triassic sediments are thick in the southern part and thin to only about 1000 m in the northern part. They are overlain by only thin Jurassic and Cretaceous strata, which were partly eroded in the southern part (Fig. 3.4.18, middle and upper part; Friberg 2001). Tertiary deposits reach more than one kilometre thickness in the southern part.

The palaeotemperature distribution was influenced to a great extent by the burial and erosion pattern as described above. Other controlling factors were basal heat flow, igneous and hydrothermal activity, and advective heat transport caused by water circulation (Magri et al. this volume). The general maturity pattern as observed in many deep wells in the CEBS is quite uniform, i.e., there is a systematic increase of maturity with depth. Extreme maturity gradients as observed in other basins which were clearly influenced by igneous activity (e.g., Sachsenhofer and Littke 1993) are absent in the CEBS, even in areas which were suspected of being influenced by igneous intrusions. An exception is the very northern part of the

Subsidence, inversion and evolution of the thermal field Figure 3.4.17. Profile C-C’ through the transition zone between Pompeckj Block in the north and Lower Saxony Basin in the south. Three time steps are shown Middle Jurassic, Late Cretaceous and presentday

basin, where small igneous dykes and sills disturb the general maturity pattern, leading to locally high vitrinite reflectance values (Friberg 2001). High thermal maturities at the Earth’s surface are found in areas, where strong inversion took place. This is the case in the Lower Saxony Basin, in parts of the Netherlands, and in the Polish Basin. For a given stratigraphic horizon, thermal maturities are mainly controlled by deepest burial during basin subsidence. In the area of the Pompeckj Block, this stage was reached late, i.e., in the Neogene. Depth of present burial does therefore correlate well with maturity there (Fig. 3.4.19). This is,

however, not the case in other areas, where deepest burial was reached sometime in the past, usually in Jurassic or Cretaceous times. Examples are the areas at the southern basin margin. Burial trends for different sub-basins of the CEBS are shown in figure 3.4.20 for (i) the top Triassic/base Jurassic and (ii) the top Carboniferous/base Permian, according to modelling results (see Rodon and Littke 2005; Senglaub et al. 2006; Schwarzer and Littke 2007; Nelskamp et al. 2007; Resak et al. 2007 for details). The burial trends were calculated for real wells, calibrated by maturity data. They show some general trends, but the wells clearly

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R. Littke · M. Scheck-Wenderoth · M.R. Brix · S. Nelskamp Figure 3.4.18. Profile D-D’ through the East German part of the Central European Basin system. Three time steps are shown: Early Cretaceous (Wealden), Cretaceous/Tertiary Boundary and present-day

do not represent the respective entire sub-basin. A common feature of all burial trends is the enhanced subsidence during the Permian and Triassic. The onset of this phase of high sedimentation rates was, however, slightly different in the sub-basins. Early subsidence in Permian (Rotliegend) times is evident in the basin centre, i.e., the Glückstadt Graben near Hamburg, as well as in Eastern Germany and in the Polish Basin, where thick, partly volcanogenic, deposits witness the break-up phase of the basin. In contrast, in the more marginal basins in the west and northwest, strong sedimentation set in later in Zechstein or Triassic (Buntsandstein) times (see Bachmann et al. this volume). Highest Triassic subsidence rates are evident for the Glückstadt Graben (Fig. 3.4.20b) as well as for the Polish Basin.

During the Late Triassic (Keuper), Early and Middle Jurassic (Liassic, Dogger), continuous sedimentation at moderate rates prevailed throughout the basin, whereas in the Late Jurassic, strong differentiation with significant erosion phases started. In the area of the Central Netherlands Basin, the Pompeckj Block, and the Glückstadt Graben, erosion of several hundred metres occurred, partly localized due to salt tectonics. Part of the eroded material was probably deposited in the Lower Saxony Basin, in which a phase of strong subsidence started in the latest Jurassic and Early Cretaceous (Fig. 3.4.20a). In the centre of the Lower Saxony Basin, up to 4 kilometres of sediments were deposited then. This period was characterised in most other areas by slow subsidence. In the Late Cretaceous, inversion affected several sub-basins, especially

Subsidence, inversion and evolution of the thermal field

Figure 3.4.19. Present-day maturity (vitrinite reflectance) map of the top Carboniferous in the CEBS (modified after Teichmüller 1984)

the Lower Saxony Basin, the West and Central Netherlands basins, and the Polish Basin. The onset of inversion and erosion has been deduced from a variety of geological factors and – most importantly – from fission track data. However, it must be noted that the inversion phase did not affect the entire CEBS: in the former (Permian) basin centre with the Pompeckj Block and the Glückstadt Graben, sedimentation resumed at that time. Finally, in the Triassic, enhanced sedimentation mainly affected the Central and West Netherlands basin, but also the Pompeckj Block and was strongly related to the development of the North Sea Central Graben (see Sirokko et al. this volume). Fission track thermochronology (see box) provides information on time slices during which samples were exposed to specific temperatures. While zircons register the temperature evolution between 180 and 380 °C, apatites are sensitive from 60 to 120 °C. If the temperature in strata being buried stays within the range of sensitivity of the considered mineral, the complete evolution from burial through inversion to uplift and cooling may be traced. If the temperature becomes too high and all tracks are annealed, then only the late cooling part of the thermal evolution is represented in the fission track record. Zircons in

deeply buried basins thus usually represent a longer and older segment of the thermal path, while apatites often cover only a relatively late part of the cooling history. The thermal record witnessed by zircon fission tracks in the Northwest German Basin starts with the deposition of sediments during Late Carboniferous times in the northern foredeep of the Variscan orogen. Carboniferous samples from the LSB reflect different degrees of annealing (Senglaub et al. 2005). Some yielded ages significantly younger than the respective sedimentation ages. Single grain ages often display a strong spread, usually with a distinct tail including old ages and a steep slope towards Permian ages. This annealing is primarily interpreted as the effect of widespread magmatism followed by the circulation of hydrothermal fluids, coinciding with rifting and crustal thinning in the Southern Permian Basin. Only a single sample from the area formerly attributed to the supposed Late Cretaceous intrusion of Bramsche yielded a sharp peak around 133±6 Ma, indicating rapid cooling after annealing of all tracks. The maximum temperature reached by this sample has been modelled by Senglaub et al. (2005) as ~350 °C. The thermal pulse may be related to extension within the LSB. Similar peaks were also ob-

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R. Littke · M. Scheck-Wenderoth · M.R. Brix · S. Nelskamp Figure 3.4.20. Burial trends for the top Triassic/base Jurassic (a) and the top Carboniferous/base Permian (b) according to modelling results for different subbasins

served in some Mesozoic samples. Most of these are rather inhomogeneous. The single grain age distributions are often bimodal with a clear tail including old ages and a steep slope towards young ages between 130 and 140 Ma. On the sample age level, however, no significant annealing is observed, indicating that the samples may have been subjected to the thermal conditions of the low temperature part of the ZPAZ for only a short period of time. As most samples had been heated to more than 150 °C, apatite fission track data only reflect the cooling after the thermal maximum. Some of them yielded ages around 80 Ma independently of their sedimentation ages. This

corresponds to the time of inversion in the LSB. Track length data combined with the ages indicate a relatively fast cooling of the basin centre compared to the basin margins. Cooling was accelerated close to prominent faults or flexure zones to rates of 1.5-4 °C/Ma. Similar data have also been reported by Leischner (1994) and Petmecky (1998) further to the west and to the east. Other samples, especially from boreholes, reflect another cooling episode in the Miocene. Although some of these samples were taken in the depth at temperatures in excess of 125 °C, Cretaceous single grain ages were observed. This indicates that the samples were only exposed to these elevated temperatures rather recently after continuous cooling during the Tertiary.

Subsidence, inversion and evolution of the thermal field Figure 3.4.21. Temperature-Time curves for the top Triassic/base Jurassic (a) and the top Carboniferous/base Permian (b) according to modelling results for different subbasins

On the regional scale, the zircon fission track ages display a decrease towards the north. The southern part of the LSB and the adjacent Münsterland block are characterized by ages essentially equivalent to depositional ages, while reset Permian ages are specific for the northern part. Only a single sample from the basin centre yielded an Early Cretaceous cooling age after complete annealing of the tracks. Apatite fission track ages also become younger from south to north within the basin. The decrease is not gradual. Abrupt changes coincide with a major flexure zone (Wiehengebirge) and indicate that the rocks south of the flexure zone have been more uplifted and eroded than those further to the

north. This is in agreement with Petmecky et al. (1999) who observed further east that the Uchte anomaly is also limited by structural elements. The interpretation of fission track and temperature related maturity data and especially their conversion to amounts or rates of burial or uplift requires information on palaeo geothermal gradients and/or heat flow. This, however, is scarce and sometimes contradictory. A heat flow of ~85 mW/m² in Late Carboniferous/Early Permian times was assumed by Büker et al. (1995) for the Münsterland Block south of the LSB. Buntebarth (1985) proposed geothermal gradients between 65 and 92 ° C/km,

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Figure 3.4.22. Maturity-Time curves for the top Triassic/base Jurassic (a) and the top Carboniferous/base Permian (b) according to modelling results for different subbasins

corresponding to heat flows of 140-175 mW/m², for the time of maximum burial of the LSB in the Early Cretaceous for empirical reasons. Leischner et al. (1993), however, assumed a heat flow of 80 mW/m² at this stage based on fission track and fluid inclusion studies as well as on maturity-vs.-depth profiles. This value is also supported by Senglaub et al. (2006), who numerically modelled heat flows of 60-70 mW/m², taking into account a calibration by multiple maturity indicators. Extending this data set by fluid inclusion data, Adriasola-Muñoz et al. (2007) revealed periods of elevated heat flow during the Early Permian (magmatism), the Liassic (basin differentiation), and the Late Jurassic/Early Cretaceous (rifting). Depending

on the position within the basin, heat flow during Tertiary times has been deduced to have been ~45 mW/m², corresponding to the lower part of the range of present-day heat flow in North Germany and the North Sea between 40 and 60 mW/m² (Hurter and Haenel 2002) or even 80 mW/m² (Leischner et al. 1993; Senglaub et al. 2006). Temperature evolution over time closely followed the burial history and is shown for the same wells/regions in figure 3.4.21. In the Glückstadt Graben, representing the basin centre, high temperatures of 250 °C for the top Carboniferous were already reached during the Jurassic, and even higher temperatures during the Tertiary. Very high

Subsidence, inversion and evolution of the thermal field

temperatures also occurred in parts of the Lower Saxony Basin during the Cretaceous, getting up to between 250 and 350 °C at the top Carboniferous. These high temperatures were lowered by about 150 °C during the inversion phase. Only in the Lower Saxony Basin the top Triassic did also reach high temperatures of 200 °C and more. In all other sub-basins, this layer usually reached temperatures of 80-100 °C. In the Pompeckj Block and Eastern Germany, highest temperatures were only reached in the Tertiary, favouring late gas generation from Palaeozoic source rocks. This is locally also the case in the Central and West Netherlands basins. In contrast, maximum temperatures in the Lower Saxony Basin and Polish Basin clearly occurred earlier, before inversion.

too high (di Primio et al. this volume). A further slight increase in maturation during the Tertiary may have led to generation of some nitrogen (Krooss et al. this volume) but no significant amounts of hydrocarbons. Another extreme case is the Lower Saxony Basin, in which maturation occurred during the Early Cretaceous, also affecting the Mesozoic rocks to such an extent that they partly reached the gas generation zone, i.e., vitrinite reflectance values of 1.3% and above (Fig. 3.4.22a). This is clearly the case for the Liassic Posidonia Shale (AdriasolaMuñoz et al. 2007) which is known as the principle oil source rock of the basin, but also for the younger Wealden Shales. A more continuous maturation characterises the other subbasins/wells presented in figure 3.4.22.

The information given above was used to calculate thermal maturity, i.e., vitrinite reflectance, for the different sub-basins (Fig. 3.4.22). The extreme case is the Permian basin centre (Glückstadt Graben) in which major maturation was clearly related to the Triassic/Jurassic period. For that area, it can be expected that virtually no methane was generated from the coal-bearing Carboniferous source rocks after the Triassic because temperatures were

In summary, detailed basin studies are essential in order to understand temperature histories and petroleum generation in a basin as complex as the CNBS, which is divided into different sub-basins. Perhaps the greatest difficulty is unravelling the temperature evolution for times of inversion and erosion. The only tools to tackle this problem are numerical models fed by a variety of geological, geochemical and geophysical parameters.

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4

Chapter 4

Basin fill

4

Chapter 4.1

Depositional history and sedimentary cycles in the Central European Basin System G.H. Bachmann · T. Voigt · U. Bayer · H. von Eynatten · B. Legler · R. Littke

4.1.1 Palaeoclimate, Palaeogeography, and Palaeoenvironment The sedimentary evolution of the Central European Basin System (CEBS) was driven by both external and internal forces. Whereas internal forces comprise tectonic processes as well as salt-driven and fluid-driven ones, external forces are related to climate, geography, environment and sea level changes during the 300 million years of basin evolution. The latter processes are treated in this part of the book and in this chapter, with special emphasis on gross depositional cycles and the origin of the sediments or specific minerals therein. Sedimentation in the CEBS responded greatly to palaeogeographic and palaeoclimatic conditions and their changes over time. Palaeobathymetry, i.e., the sedimentation depth below sea level, is often difficult to evaluate, even when fossils such as benthic foraminifers can give important intimations. Other evidence is derived from the

sediments, e.g., coals usually witness terrestrial conditions, whereas reefs and shore face sands indicate shallow water conditions. Sea surface temperatures are often derived from oxygen isotope data measured both on planktonic and benthonic forams (e.g., Pearson et al. 2001), and also on specific organic molecules such as unsaturated ketones (Müller et al. 1998), which respond in their structure to sea surface temperature changes. Wygrala (1989) summarised surface temperature trends as a time-latitude diagram (Fig. 4.1.1) which provides useful clues to the evolution of surface temperature: the palaeolatitude position of the CEBS is plotted, so that the general trend for this area becomes visible. Roughly, it is well established that Late Carboniferous sedimentation of coal-bearing strata took place close to the equator under tropical, humid conditions. A drop in global temperature in the latest Carboniferous/earliest Permian and northward movement of the study area led to more arid conditions during the Permian and Early Triassic. This is witnessed by both red coloured clastic sediments and salts (Warren this volume). Sedimentation took place either under terrestrial

Figure 4.1.1. Surface temperature versus latitude versus geological time based on Wygrala (1989). The pathway of a continental fragment from Central Europe is shown

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Figure 4.1.2. Palaeogeographic overview: A – Late Permian: Upper Rotliegend II (Ziegler 1990; Coward et al. 2003); B – Late Triassic: Weser Fm. (Ziegler 1990; Fisher and Mudge 1998; Dadlez et al. 1998a; Barnasch and Franz pers. comm.); C – Late Jurassic: Oxfordian (IGCP86 1987; Ziegler 1990; Coward et al. 2003); D – Late Cretaceous: Cenomanian/Turonian (Hancock and Scholle 1975; Ziegler 1990; Coward et al. 2003); E – Palaeogene: Palaeocene (Coward et al. 2003). Note that some “areas of non-deposition/erosion” received sediments at the respective times, which were, however, later eroded

Depositional history and sedimentary cycles in the Central European Basin System

conditions or in shallow water. The same holds true for the Middle and Late Triassic. The latter period is characterised by more abundant higher plant fragments (coaly organic matter) again indicating a more humid climate. In contrast to the Triassic, the Jurassic sedimentation in the CEBS took place under predominately marine conditions. However, deep water conditions were not established, as evident from the occurrence of shallow water sandstones and limestones. During this period, but even more during the Cretaceous greenhouse climate, high surface temperatures can be expected due to the palaeogeographic position and palaeotemperatures (Fig. 4.1.1). A change in the arid climatic conditions in the Late Jurassic is reflected by the deposition of evaporites, stromatolite-bearing marlstones and in general, the predominance of limestone deposition. Further sedimentation shifted from terrestrial-shallow marine in the earliest Cretaceous (“Wealden”) towards deeper marine conditions in the Albian and Late Cretaceous and back to shallow marine/terrestrial in the Tertiary. The last phase of high surface temperatures occurred in the Eocene; thereafter northward movement of the area and a drop in global temperatures led to the present-day situation. During the Pleistocene, the Scandinavian ice sheet covered large parts of the basin several times and left behind thick glacial deposits. A more detailed view of the evolution of this area can be derived from palaeogeographic maps. Whereas the gross palaeogeographic situation of the basin can be deduced from world maps (Maystrenko et al. this volume; after Scotese 2004), a more detailed view is provided here (Figs 4.1.2, A-E). Five time slices were chosen to demonstrate the evolution of the Central European Basin System (CEBS) and its surrounding areas from the Late Palaeozoic to the Cenozoic. The maps, as well as the short description, are mainly based on Ziegler (1990) and Coward et al. (2003). The basin evolution started in the latest Carboniferous to Early Permian with intensive and widespread volcanism. Lower Rotliegend and Upper Rotliegend I-sedimentation was restricted to small pull-apart basins. Further extension and the beginning of regional thermal subsidence resulted in the formation of two large depositional basins, the Northern and Southern Permian Basin (NPB, SPB), which were filled with Upper Rotliegend II-sediments (Fig. 4.1.2 A). The SPB reached from England to Poland and the NPB from the Northern North Sea to Denmark. They were separated from each other by an alignment of highs (Mid North Sea High, Ringkøbing-Fyn-Møn High). Continental siliciclastics were deposited in a desert environment. Lacustrine evaporites were formed in the basin centre of the SPB. Both basins were flooded by the Zechstein transgression, which resulted in the deposition of a thick series of carbonates and evaporites. The basin

system was connected towards the Boreal Ocean via the Viking Graben between Fennoscandia and the Shetland Platform but a faunal exchange with Tethys also existed. The highs, separating the two large depositional basins during the Permian, were gradually included in the depositional area during the Triassic. The Early Triassic was characterised by a regression and the recurrence of continental deposition as well as by the onset of rifting processes, which triggered deposition. Siliciclastic sedimentation of the Early Triassic Buntsandstein was influenced by marine ingressions from the southeast (East Carpathian and Silesian-Moravian Gateways). During the Middle Triassic a pronounced transgression from the southeast and later also from the southwest (AllemannicBurgundy Gateway) resulted in the deposition of calcareous Muschelkalk deposits. Continental deposition persisted in the central and northern North Sea area. The Late Triassic was characterised by continuing sedimentation of continental siliciclastics and in parts by deposition of evaporites (Fig. 4.1.2 B, Weser Fm.). Marine ingressions from the Tethyan realm occurred mainly via the Allemannic-Burgundy Gateway. With a renewed transgression, marine deposition recurred at the beginning of the Jurassic and persisted until the Cenozoic. A communication from the Tethys to the Arctic seas was established during the Early Jurassic and shallow marine shales were deposited in epi-continental seas. During the Middle Jurassic, the basin was restructured by the onset of North Sea doming as a result of intensive rifting and crustal separation in the Central Atlantic. Rifting established a system of deeply subsided graben structures and related strike-slip faults across the CEBS (e.g., Lower Saxony Basin, Central Netherland Basin) which accommodated several thousand metres of sediment. Uplifted graben shoulders contributed clastic material to the basin fill. Rotation of blocks at listric normal faults caused unconformities to develop. Deposition of continental clastics took place in grabens in the central and northern North Sea. South and north of the uplifted area, marine siliciclastics and volcanics dominated. During the Late Jurassic the Mid North Sea High was drowned in its northern parts but areas at its southern vicinity fell dry. Deposition (for example the Oxfordian, Fig. 4.1.2 C) was characterised by marine shales in the northern and western North Sea as well as in wide areas of Britain. Carbonates were deposited from southern England across the Lower Saxony Basin to Poland and at the Franconian Platform to the South. The tectonic situation was dominated by the North Atlantic rifting during the Early Cretaceous. Deposition was influenced by sea level fluctuations and Jurassic carbonate sedimentation was replaced by siliciclastics. During the

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Figure 4.1.3. Major Depositional Cycles 1–9 (2nd order), major unconformities, peak transgressions, rifting phases and halotectonic phases of the Central European Basin System. Upper, regressive, phases of DC4–9 seem to correspond with rifting and halotectonic phases. Chronostratigraphy and numerical ages after DSK (2002), Bachmann and Kozur (2004), Menning and Hendrich (2005) and Kozur and Weems (2007). Z = Zechstein, B = Buntsandstein, M = Muschelkalk; Unconformities: VU = Variscan, SU = Saalian, A I = Altmark I, A III = Altmark III, HU = Hardegsen (base Solling Fm.), ECU = Early Cimmerian (base Arnstadt Fm.), MCU = Middle Cimmerian (base Dogger Sandstones), LCU = Late Cimmerian (base Bückeberg Fm.), LU = Laramide (Base Palaeocene). Star: major volcanism. Further interpretation see Fig. 4.1.4

Depositional history and sedimentary cycles in the Central European Basin System

Late Cretaceous the CEBS was affected by a major sea level rise, basin margins were overstepped and deposition was dominated by marls and chalk from the Cenomanian until the Danian (Fig. 4.1.2 D). Siliciclastic input was restricted to Fennoscandia and the Rhenish-Bohemian-Silesian massifs during the Cenomanian/Turonian. In contrast to the Early Cretaceous, the Late Cretaceous and Palaeocene reflected the opening of the North Atlantic and the northward drift of the African plate. This fundamental change in the tectonic regime resulted in a change of stress fields in central and northwest Europe. Compressional stresses caused the inversion of parts of the CEBS during Late Cretaceous. Inverted structures shed progressively clastic material into the chalk sea. Low global sea level, regional uplift of the southern basin margin of the CEBS and especially of the thermally uplifted Shetland Scottish highs resulted in clastic input into the North Sea Basin during the Palaeocene (Fig. 4.2.1 E). The opening of oceanic basins in the North Atlantic north of Scotland together with hot-spot activity caused volcanism until the early Eocene. Since the late Eocene, the Rhine and Eger rift systems have developed, the clastic input from the Scottish Highlands has decreased and the Polish platform has been gradually included in the depositional area. Regional uplift occurred at the southern basin margin.

4.1.2 Sedimentary cycles The basin fill of the CEBS is sometimes viewed as a mirror of sea level cycles. During the Permo-Carboniferous vast masses of inland ice existed on the southern continents and global sea level was unusually low. Sea level generally rose with the break-up of Pangaea to its Phanerozoic maximum in Late Cretaceous times of some 250 m above the Permian (and present day) level as indicated by the maximum extent of open-marine facies on all continents. This all-time sea level high was followed by a general drop of sea level to its present low due to the renewed gathering of the continents, lower sea floor spreading rates and again vast accumulations of ice in the polar regions. The scientific community was indeed surprised when an EXXON working group published a global Phanerozoic sea level record showing more than 50 rapid sea level falls (Payton 1977). The initial approach of translating coastal onlap-offlap patterns directly into sea level changes was soon heavily criticised (e.g., Hallam 1978, 1981). Indeed, nowadays we are aware that onlap-offlap or transgressionregression cycles are affected by quite a number of factors like eustatic changes, subsidence/uplift rates, and climatic variations causing variations in sediment supply. A recent review concerning the complexity of the translation of

onlap-offlap patterns into sea level changes is given by Kenneth et al. (2005). Using back-stripping techniques, they conclude that the original estimates of sea level fluctuations, including the long-term trend, are too high by a factor of approximately 2.5. Regionally observable are onlap-offlap patterns. Locally, in a well or a relatively short seismic line, however, regressive-transgressive cycles are easily established which again may not directly coincide with onlap or sea level changes. The term transgressive-regressive cycle will, therefore, be used here. The long-term global transgressive-regressive cycle from the Permo-Carboniferous to the present has been referred to as a “1st Order Cycle” or “Continental Encroachment Cycle” and was the second such in the Phanerozoic eon (Vail et al. 1977; Fischer 1984; Duval et al. 1998). The Continental Encroachment Cycle can be subdivided into a number of transgressiveregressive 2nd Order Cycles, each normally lasting between 3 and 50 million years (Ma). They roughly correspond to the transgressive-regressive sequences of Sloss (1963) or the major transgressive-regressive cycles of Jacquin and Graciansky (1998) and of other authors (although they are often difficult to correlate). Second order cycles were caused by changes in basin subsidence and eustatic sea level. 3rd Order Cycles (0.5 to 3 Ma) correspond to the sequences of sequence stratigraphic or base level concepts (e. g., Vail et al. 1977; Haq et al. 1988) and were caused by changes in basin subsidence and eustatic sea level. 3rd order cycles consist, in turn, of 4th and 5th Order Cycles (0.01 to 0.5 Ma), which correspond to parasequences, small base-level cycles or astronomically controlled Milankovitch cycles of the order of ~0.02, 0.04, 0.1 and 0.4 Ma (Strasser et al. 2006). Third order or base-level cycles were recognised in the marginal seas of the CEBS as far back as 1917 by Klüpfel. These coarsening-upward cycles are typically terminated by a condensed, reworked fossil-rich layer in near coastal environments and are frequently associated with changes in the fossil assemblages over time, with some indications of adaptation as well as faunal overturn by immigrants. Both aspects, sedimentary and evolutionary cycles, are discussed by a series of authors in Bayer and Seilacher (1985) with special focus on the marginal seas of the CEBS. The following overview aims to outline the major trends of the CEBS fill by defining and describing the principal depositional cycles, which are for practical reasons referred to as “Depositional cycles (DC)”. Most of the following description and interpretation is based on work in the Central (German) part of the CEBS but it applies in part to other areas as well. Earlier definitions of major depositional cycles were given by Jacquin and Graciansky

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(1998) and Robin et al. (2000) in more marginal parts of the CEBS. Nine important depositional cycles (DC 1-9) have been observed ranging from a few Ma up to 75 Ma (Fig. 4.1.3). Each cycle is bounded by major erosional unconformities. The first depositional cycle (DC 1) consists of volcanic and some terrestrial clastic deposits. The second and third cycles (DC 2-3) are characterised by terrestrial clastic sedimentation of fining-upward character. Each of the succeeding six cycles (DC 4-9) shows a transgressive phase, peak transgression and a regressive phase. Instead of sea level changes, the cycles can be better described in terms of rising and falling base levels (Fig. 4.1.4), which seems especially appropriate for DC 1-3 which were formed in a period when continental depositional systems were only temporarily linked to the global sea level. The short duration of the first depositional cycles is due to the fact that early stages of the evolution of the CEBS were strongly triggered by short-term tectonic events and not so much by global sea level changes. Furthermore, these stages were studied in great detail as they represent the main reservoirs of hydrocarbons in the basin. It seems that the upper, regressive, phases of DC 4-8 are closely linked to major tectonic and halotectonic events. However, the stratigraphic record of the CEBS is rather inhomogeneous, partly based on lithostratigraphic – facies based – units, partly on well defined chronostratic ones. Subdivision in terms of formations lacking detailed time control is especially due for the Permian and Triassic, while Jurassic and Cretaceous are in general excellently bounded chronostratigraphically. Observed or interpreted cyclicity, therefore, may be essentially hampered by precise chronostratic data not only in time but also in space. The transgressiv-regressive cycles given here are indeed preliminary, needing partly chronostratigraphic control and/or further correlations with areas providing chronostratigaphic control.

4.1.2.1 Depositional cycle 1: Altmark (latest Carboniferous to Early Permian) Variscan compressional tectonics came to an end in Late Carboniferous times. This was followed by regional uplift and erosion. Initiation of the CEBS was accompanied by strong magmatism that can be related to strike-slip faulting and lithospheric thinning (Arthaud and Matte 1977; Bachmann and Grosse 1989; Ziegler 1990; Bachmann and Hoffmann 1995, 1997). Most faults trend NE-SW and WNW-ESE. The CEBS evolved on the former Variscan foreland but its southern parts encroached into the successively peneplained Variscan fold-belt as well (Maystrenko

et al. this volume). The Altmark Subgroup and equivalents (Stephanian/Ghezelian–Asselian) are here considered as the oldest deposit of the CEBS, at least in northern Germany (Plein 1995). The succession consists of multiphase acidic and intermediate volcanics as well as some redbeds of up to 2-3 km thickness (Geluk 2005).

4.1.2.2 Depositional cycle 2: Müritz (Early Permian) On top of the Altmark Subgroup and equivalents, as well as on older sediments, another unconformity developed (“Saalian Unconformity”, Stille 1924; Ziegler 1990) that was interpreted to be the result of regional thermal uplift in the CEBS area due to crustal heating (Bachmann and Hoffmann 1997). The irregular topography of volcanic extrusives may also play a role in parts of the basin (Geißler et al. 2008; Breitkreuz et al. this volume). On top the Müritz Subgroup, a succession of red and grey to black fluvio-lacustrine sediments follows. They show an overall fining-upward pattern. The Müritz Subgroup occurs mainly in a WNW-ESE trending area between NE Germany and the German North Sea with depocentres in several interconnected transtensional rift basins (Plein 1995; Geluk 2005). Based on K-Ar ages, contemporaneous extrusive volcanic activity may have been restricted to the Central North Sea Graben and Horn Graben (Geluk 2005).

4.1.2.3 Depositional cycle 3: Havel (Middle Permian) The “Altmark I” unconformity (Hoffmann et al. 1989) on top of the Müritz Subgroup represents another long period of erosion in parts of the CEBS, lasting approximately 15 Ma (late Early Permian to early Middle Permian). Local basalts may record a generally extensional regime (Hoffmann et al. 1989). The main trend of the CEBS was now WNW-ESE. The Havel Subgroup above the unconformity can be more than 1100 m thick in northern Germany (Plein 1995). It consists of the Parchim and Mirow formations. They are separated by the “Altmark II” unconformity, which is especially pronounced at the basin margins. Both formations consist of predominantly red alluvial, fluvial, playa and salt lake sediments of generally fining-upward character. Initial deposition took place in rift basins between the German North Sea and Poland (Geluk 2005). In parts of North Germany, at the southern basin margin, the equivalents of both formations are confined to a N-S trending rift system (Gast 1988). DC 3 shows rapid subsidence due to rifting and the beginning of thermal relaxation (thermal subsidence).

Depositional history and sedimentary cycles in the Central European Basin System

Figure 4.1.4. Scheme of a transgressive-regressive depositional cycle (2nd order in Fig. 4.1.3) with associated rifting and halotectonic phases. Alternative interpretations as baselevel cycle or sequence

4.1.2.4 Depositional cycle 4: Elbe, Zechstein, Lower and Middle Buntsandstein parts (Late Permian to Early Triassic) The base DC 4 is marked by the “Altmark III” unconformity, which is especially distinct at the basin margins. DC 4 is characterised by rapid subsidence and the expansion of the CEBS into eastern England and eastern Poland. The basin had three depocentres: the AngloDutch, the North German and the Polish Basin (Geluk 2005). A number of rather conservative NNE trending swells and rifts (“troughs”) differentiated the generally W-E trending CEBS, including the Netherlands Swell, Ems Trough, Hunte Swell, Weser Trough and EichsfeldAltmark Swell. It seems that the swells coincide with the rift-shoulders of the troughs (Frisch and Kockel 1997). As sedimentation did not keep pace with the subsidence rates of the CEB its surface was temporarily below sea level (Glennie and Buller 1983). The Dethlingen and Hannover formations of the Elbe Subgroup consist of predominantly red alluvial, fluvial, aeolian, playa and salt lake sediments that form a fining-upward succession of up to 1500 m thickness. The cyclicity of the sandstones, shales and halites has been interpreted as ~400,000 and ~100,000 years Milankovitch eccentricity cycles (Gast 1995). First minor marine ingressions occurred in the Dethlingen Formation and then again in the uppermost Hannover Formation possibly through the proto-Viking and proto-Horn grabens to the north (Fig. 4.1.2). At the beginning of the Zechstein Group the basin was flooded with sea water (peak transgression). This basinwide transgression led to the deposition of the euxinic, starved, thin “Kupferschiefer” (Paul 2006a). This was the first widespread marine ingression since the Westphalian Aegir ingression which occured some 50 Ma earlier during the Late Carboniferous.

The original thickness of the Zechstein Group is difficult to ascertain due to later halotectonics (Scheck-Wenderoth et al. this volume), but has been at least 1500 m in some areas. The Zechstein consists of several evaporation cycles. The individual cycles, if ideally developed, exhibit successions of shales, carbonates, sulphates, and chlorides (Warren this volume). However, the Zechstein as a whole shows a generally regressive and hypersaline trend. Only the lower 3 cycles have the full cyclic inventory including marine shales and carbonates; the upper 4 cycles are dominated by shales, sulphates and halites and become more and more confined to the central parts of the CEBS. In the marginal parts they are represented by red brittle evaporitic shales of the “Bröckelschiefer facies” that were deposited in a sabkha environment. Zechstein sequence stratigraphy was outlined by Strohmenger et al. (1996). The succeeding Lower Buntsandstein Subgroup is up to 400 m thick and comprises the Calvörde and Bernburg formations. Both formations consist of mostly red shales, sandstones and oolithic and stromatolitic limestones without evaporites, and were deposited in fluvial, lacustrine and playa environments, respectively. In the eastern and central parts of the CEBS, however, microfossils (acritarchs, brackish ostracods and foraminifers) indicate some brackish influence (Pienkowski 1991). The Zechstein/Buntsandstein boundary is characterised by the influx of more sandy material (Best 1989) due to higher rates of precipitation associated with the Permian/Triassic boundary event. The biostratigraphic Permian/Triassic boundary is some 20 m above the lithostratigraphic Zechstein/Buntsandstein boundary (Kozur 1999). The Middle Buntsandstein of DC 5 comprises the Volpriehausen, Detfurth, and Hardegsen formations consisting of red sandstones, shales and thin oolites deposited in fluvial, lacustrine, and playa environments, respectively. Minor marine to brackish influence is recorded in lower Middle Buntsandstein deposits of Germany and

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Poland (Wycisk 1984; Becker 2005). Connection to the Tethys through eastern Poland is generally favoured and is compatible with the westward decrease of marine influence in the Lower and Middle Buntsandstein. The Middle Buntsandstein is normally a few 100 meters thick except in the Horn Graben (3500 m), the Glückstadt Graben (2000 m), the Mid-Polish Trough (more than 1000 m) and the Central Graben (500 m). Short tectonic uplift pulses have been identified at particular structures within the Middle Buntsandstein (pre-Volpriehausen, pre-Detfurth, preSolling and intra-Solling). Of these, the pre-Solling pulse, which gave rise to the “Hardegsen” unconformity, was the strongest (Röhling 1991) and terminates DC 4. Buntsandstein sequence stratigraphy was outlined by Aigner and Bachmann (1992). The small-scale cyclicity of the Lower and Middle Buntsandstein is interpreted as ~100,000 years Milankovitch eccentricity cycles (Geluk and Röhling 1999; Bachmann and Kozur 2004).

4.1.2.5 Depositional cycle 5: Middle Buntsandstein part (Early Triassic) to Middle Keuper part (Carnian) The base DC 5 is bounded by the prominent late Early Triassic (Olenekian) “Hardegsen Unconformity”. DC 5 begins with mostly coarse-grained and thin fluvial sandstones of the uppermost Middle Buntsandstein (Solling Formation). The uppermost Solling Formation and the Upper Buntsandstein (Röt Formation) are characterised by the onset of a marine ingression. Halite and gypsum deposits are common in the lower Röt Formation and mostly fine-grained red clastics in its upper part. The main transgression began during the Anisian (uppermost Röt) and a shallow sea with carbonate sedimentation (Myophoria beds) spread gradually from Poland into eastern Germany (Kedzierski 2000); this heralded the Muschelkalk transgression which subsequently flooded most of the CEBS. The carbonate-dominated Muschelkalk Group (Anisian and early Ladinian) is divided into three subgroups. In average up to 400 m of Muschelkalk were deposited, but the combined thickness of the Muschelkalk and Röt is up to 3000 m in the Horn Graben and 4000 m in the Glückstadt Graben (Baldschuhn et al. 2001). The Middle Muschelkalk is characterised by evaporites. Unusually thick Middle Muschelkalk in the above mentioned grabens is mainly due to thick halite deposits. Renewed opening of the eastern gateways and the new Allemannic-Burgundy gateway to the Tethys allowed the establishment of a large carbonate ramp system in the CEBS during the Upper Muschelkalk. Peak transgression of DC 5 occurred in the latest Anisian part of the Meissner Formation (Aigner and Bachmann 1992). The upper, early Ladinian, part of the Upper Muschelkalk is generally regressive. In the north-

eastern CEBS it grades into reddish and greyish marls, sandstones and dolomites deposited under brackish conditions. During the late Ladinian, the marine conditions of the Muschelkalk Group were rapidly replaced by deltaic, hypersaline and non-marine environments of the Keuper Group. The Keuper is typically 400–600 m thick in most parts of the CEBS, but may increase to more than 2000 m in syndepositional grabens, with up to 6000 m present in the Glückstadt Graben (Maystrenko et al. 2005a). The Keuper Group is subdivided into three subgroups (Stollhofen et al. this volume). The Lower Keuper (Erfurt Formation) consists of predominantly grey shales, sandstones and dolomites that accumulated in environments ranging from restricted marine to lacustrine. Its thickness is rather uniform (80-125 m), except in the Glückstadt Graben (600 m), the Ems Low, the Lodz Graben and the Dutch Central Graben (200 m; Geluk 2005). In contrast, the varicoloured shales, evaporites and carbonates of the lower Middle Keuper (Grabfeld Formation) were largely deposited in sabkha and playa systems; pedogenic features are common. Brief marine ingressions occurred in parts of the CEBS during sea level highstands and deposited thin carbonate beds with impoverished marine faunas. Extensional tectonics led to a large number of fault-bounded depressions in which 1000-2000 m and more of lower Middle Keuper were deposited containing thick halites. The Stuttgart Formation of the Middle Keuper is between 20 m to more than 100 m thick and comprises mudstones and channelised sandstones deposited in large fluvio-deltaic systems that spread from uplifted northwestern Scandinavia over the CEBS towards the Tethys (Wurster 1964; Kozur and Weems 2007; Paul et al. in press). The Weser Formation was again deposited in sabkha and playa environments and consists mostly of red shales with abundant pedogenic features and evaporites (Fig. 4.1.2 B). The formation terminates in large parts of the CEBS with several meters of stacked gypcretes, the Heldburg Gypsum, that indicates decreasing accommodation potential and is succeeded by the “Early Cimmerian” unconformity. Several unconformities have been identified for the regressive phase. The minor pre-Lower Keuper unconformity is only distinct in the southern parts of the basin. Three widespread mid-Carnian unconformities associated with the Stuttgart Formation were most likely caused by eustatic drops of sea level. The early Norian Early Cimmerian Unconformity is especially significant on swells, like the Eichsfeld-Altmark Swell (Barnach et al. 2005), and at the basin margins. Sequence stratigraphy was worked out by Aigner and Bachmann (1992; Triassic), Szulc (1999, 2000; Muschelkalk) and Nitsch (in DSK 2005; Keuper).

Depositional history and sedimentary cycles in the Central European Basin System

Box 4.1.1 Sedimentation and erosion rates In the complex Central European Basin system (CEBS) accumulation rates of sediments differed significantly throughout time. Sedimentation was interrupted by several long-lasting periods of erosion, but these periods did not affect all parts of the basin to the same extent. It is not easy to determine either the exact timing and rates of erosion, or the sedimentation rates of now-eroded sedimentary sequences. For major erosional phases such a quantification can be achieved by combining maturity information, fission track data and numerical modelling (de Jager 2003; Littke et al. this volume). An example is the Late Cretaceous erosion phase that affected the Lower Saxony Basin (LSB), the basins in the Netherlands WNB and CNB), the North-East German Basin (NEGB) and the Mid-Polish Trough (MPT). In the LSB, total erosion reached 6000 m (Petmecky et al. 1999; Senglaub et al. 2005, 2006); accordingly thermally highly mature and low permeable rocks occur at or close to the Earth surface there. Even if the absolute thickness of sedimented and later eroded rocks can be quantified, it still remains difficult to derive the exact ages of sedimentation and the erosion rates if long periods exist which are no longer documented by sediments. Whereas eroded thicknesses related to major erosional phases can be quantified, this is hardly possible for minor phases of erosion. Clearly, tectonic and sedimentary evolution cannot be understood if phases of erosion are not considered. Furthermore, hydrocarbon generation and accumulation took place to a great extent during these phases.

Figure 1: Average sedimentation and erosion rates in different parts of the Central European Basin system given in m/kyr (kyr: 1000 years). Abbreviations: ER: erosion rate, SR: sedimentation rate, WNB: West Netherlands Basin, CNB: Central Netherlands Basin, LSBwest: western Lower Saxony Basin, PB: Pompeckj Block, GG: Glückstadt Graben, HG: Horn Graben, NEGB: North East German Basin, MPT: Mid Polish Trough (Petmecky 1998; Friberg 2001; de Jager 2003; Rodon and Littke 2005; Senglaub et al. 2006; Beha et al. 2007; Nelskamp et al. 2007; Resak et al. 2007; Schwarzer and Littke 2007)

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For nine key wells in the CEBS, results on sedimentation and erosion are summarised and compared in figure 1. The Pompeckj Block (PB) situated in the central part of the basin may be regarded as a reference location for this comparison. There, thick Late Carboniferous sediments witness the pre-rift evolution of the area, which then was the foreland (molasse) basin of the rising Variscian mountain belt in the south. After a rather short and ill-defined period of erosion, Permian and Triassic sedimentation set in at average sedimentation rates which did not exceed 100 m/million years. (Clearly, short term fluctuations will have occurred which are not recorded in the average rates presented here). Over the course of the Triassic and Jurassic, sedimentation rates declined. This feature is commonly attributed to decreasing effects of the cooling of the lithosphere after the initial heating during the rifting phase. In the Late Jurassic, significant uplift and erosion occurred which was followed by renewed subsidence with moderate sedimentation rates in the Late Cretaceous and Tertiary. In comparison to this reference location, the other locations show both similarities and differences. A common feature is the Late Carboniferous foreland basin sedimentation followed by a Variscian consolidation phase and high to moderate sedimentation rates during the Permo-Triassic. Sedimentation rates commonly decrease in the course of the Triassic and are low in the Jurassic. Very high sedimentation rates existed during the Triassic in the Glückstadt Graben (GG) and quite high sedimentation rates during the Jurassic characterise the NEGB, and also the basins in the Netherlands and the LSB. Central parts of the CEBS, such as PB, MPT and GG were inverted during the Late Jurassic/Early Cretaceous, when enhanced sedimentation set in at marginal positions such as in the LSB, WNB, and CNB. The latter basins as well as the MPT and NEGB were strongly inverted during the Late Cretaceous, but the exact timing of the inversion is still ill-defined due to limited fission track data (see Littke et al, this volume; Voigt et al. this volume).

Small-scale sedimentary cycles are mostly interpreted as ~100,000 years eccentricity cycles (Bachmann and Kozur 2004; Menning et al. 2005). DC 5 corresponds with the Triassic second order transgressive-regressive cycle of Aigner and Bachmann (1992) that has, however, a different upper boundary at the base Rhaetian corresponding to first marine ingressions in the CEBS. A gradual base-level rise starts, however, above the Early Cimmerian unconformity. DC 5 also approximately corresponds to the “Eastern Tethys Cycle” of Jacquin and Graciansky (1998) that is bounded by the late Scythian (Olenekian) equivalents of the “Hardegsen” unconformity at the base, but by a different late Norian unconformity at the top.

4.1.2.6 Depositional cycle 6: Middle Keuper (Norian) to Dogger (Bajocian) The base of DC 6 is bounded by the early Norian “Early Cimmerian” unconformity (Beutler and Schüler 1978). The upper Middle Keuper (Arnstadt Formation) covers the palaeorelief of the “Early Kimmerian Unconformity” in a sheet-like manner (Barnasch et al. 2007). The Arnstadt Formation was deposited in a playa system (Reinhardt and Ricken 2000). Cyclic freshwater input is indicated by fluvial sandstones, which prograded from the Bohemian Massif over much of southern Germany. The average thickness of the Arnstadt Formation is some 100 m, being 400 m in the Glückstadt Graben. The formation has the lowest subsidence rates of the Triassic. Pedogenic dolocretes are common features.

The Upper Keuper (Exter Formation) represents the youngest (Rhaetian) part of the Triassic. In large parts of the CEBS it comprises the transition from Late Triassic nonmarine environments, through paralic systems, to marine conditions in the Early Jurassic that were caused by a generally rising sea level. The Upper Keuper is up to 300 m thick in most of the CEBS, up to 600 m thick in the Glückstadt Graben (Baldschuhn et al. 2001) and also locally thickens in other grabens in NW Germany. The Early Jurassic (Liassic) consists of dark grey, more or less marly marine shales with thin limestone beds. These interfinger to the east with shallow marine and paralic sandstones (Brand and Hoffmann 1963). The early Toarcian Posidonia Shale Formation consists of dark grey, laminated, bituminous marlstones that were deposited in a euxinic starved environment. They represent peak transgression during a sea level highstand, correspond with a global Oceanic Anoxic Event (Jenkyns 1988), and are one of the principal oil source rocks of the basin (Gaupp et al. this volume; di Primio et al. this volume). The Middle Jurassic (Dogger) is initially characterised by grey marine shales interfingering in the northeast with paralic sandstones. The palaeogeographic situation changed substantially after the late Aalenian due to the development of the North Sea rift system and an associated large rift dome. This coincides with the onset of ocean floor spreading in the Central Atlantic (Ziegler 1990). In late Aalenian and early Bajocian times sands were transported southwards from the eastern flank of the dome into the CEBS forming progradational paralic and deltaic units. Two significant unconformities are present (base late Toarcian, base late

Depositional history and sedimentary cycles in the Central European Basin System

Bajocian), which are also referred to as “Middle Cimmeran” unconformities (Ziegler 1990). DC 6 corresponds roughly to the “Ligurian Cycle” of Jacquin and Graciansky (1998) which is, however, bounded by somewhat different late Norian and late Aalenian unconformities. 3rd order cycles are widespread in the Early Jurassic and the Aalenian to Bajocian within the marginal seas of the CEBS and can be correlated from Britain to south Germany. Mostly they are of the “Klüpfel”type as detailed in Bayer and Seilacher (1985).

4.1.2.7 Depositional cycle 7: Dogger (Bajocian) to Lower Cretaceous (Berriasian) The widespread late Bajocian unconformity (E. Brand, pers. comm.), also referred to as the “Late Cimmerian” unconformity, is thought to represent the basal unconformity of DC 7. Late Middle Jurassic (late Bajocian, Callovian) to early Late Jurassic (Oxfordian to middle Kimmeridgian) times were characterised by a long-term eustatic sea level rise causing predominantly shallow marine environments in the CEBS. Bioclastic sandstones grade into Callovian shales, spiculitic sandy limestones (Oxfordian/Heersum Formation), oolitic limestones (Korallenoolith Formation) and, finally, into limestones and marls (“North German Kimmeridge”/ Süntel Formation). Peak transgression is in the limestones and marls of the middle Süntel Formation (middle Kimmeridgian, Aulacostephanus eudoxus Zone; Jacquin and Graciansky 1998) and was followed by deposition of latest Kimmerian to earliest Tithonian marine limestones (upper Süntel Formation, Gigaskalk Formation). These rapidly gave way to mixed marine, brackish and limnic limestones and shales (Eimbeckhausen Formation). The latest Tithonian to earliest Berriasian Münder Formation consists mainly of grey marlstones with anhydrites and halites. Thicknesses are up to 2000 m in the Lower Saxony Basin (Senglaub et al. 2006), but vary highly because of intense syndepositional rifting and erosion of part of this layer during the Late Cretaceous. DC 7 corresponds to the “North Sea Cycle” of Jacquin and Graciansky (1998) which is, however, bounded by somewhat different late Aalenian and late Berriasian boundaries.

4.1.2.8 Depositional cycle 8: Cretaceous The widespread mid-Berriasian unconformity between the Serpulit and Bückeberg formations (“North German Wealden”), which is also referred to as the “Late Cimme-

rian” unconformity, (Ziegler 1990) can be regarded as the basal unconformity of DC 8. Continental deposits of the Early Cretaceous Bückeberg Formation spread across the basal unconformity and overlie strata on the Pompeckj Swell that are much older than Late Jurassic. The coal bearing fluvial and coastal plain sand- and mudstones of the “Wealden facies” are up to 500 m thick. A major marine transgression marks the onset of the Valanginian in the rapidly expanding Lower Saxony Basin. The longterm eustatic sea level rise in the Valanginian to Albian led to deposition of up to 3000 m relatively uniform dark grey marine shales that finally spread from England to Poland. Sandstones occur at the basin margins. At the Cenomanian boundary these shales pass into light grey pelagic coccolith marl- and limestones in the course of a major sea level rise which led to the flooding of most of the Rhenish and Bohemian Massifs. Eustatic sea level rise culminated either in Turonian (e.g., Vail et al. 1977) or Campanian times (Hancock 1989) at the worldwide Phanerozoic sea level maximum of some 250 m above the Permian (and present-day) level; this corresponds with the maximum landward extent of open-marine facies (Hancock and Kauffmann 1979). Around the Cenomanian/Turonian boundary there are several layers of black, laminated, euxinic marlstones that are interbedded in light grey marl- and limestones and were caused by a global Oceanic Anoxic Event (OAE, Schlanger and Jenkyns 1976). Reddish marlstone- limestone alternations are characteristic of early- to middle Turonian well-oxygenated shallow water facies (“Rotpläner”, Söhlde Formation). The Late Cretaceous of the basin centre is generally characterised by the deposition of widespread coccolithic marlstones and chalk that persisted to the earliest Tertiary (Danian). Especially the southern part of the CEBS and the northern boundary of the CEBS were uplifted in Late Cretaceous times (Ziegler 1990). The Late Cretaceous compressional tectonics which affected the whole of western and central Europe are also referred to as the“Subhercynian tectonic event” (Baldschuhn et al. 1991; Mortimore et al. 1998; Voigt et al. this volume). Large basement blocks like the Harz Mountains, the Flechtingen and Lausitz blocks were thrust up several 1000 m on mostly NW-SE running reverse faults. Former graben structures, like the Lower Saxony Basin and the Central Netherlands Basin, became inverted. The uplifted areas were fringed with littoral sandstones and conglomerates. Commonly, a major hiatus (time of non-deposition) developed above the “Subhercynian Unconformity”. In the southern basin, renewed sedimentation started in the Eocene above a peneplain cutting through inverted structures and deformed Late Cretaceous deposits. Campanian deposits were partly involved in deformation but may also

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Box 4.1.2 How to analyse the provenance of clastic sediment? Provenance analysis of clastic sediments aims at unravelling the pre-depositional history of the clastic material deposited in a sedimentary basin. Information on sediment provenance is crucial for, e.g., understanding the parent-rock assemblages and tectonic setting of the source area, weathering, erosion, and the linkage between source and sink (transport system). Tracing provenance changes through time and space enables us to reconstruct the major geological processes that control earth history. Given a high-resolution stratigraphic record, detailed provenance analysis provides a unique tool to derive an excellent record of tectonic processes in source areas, basin configuration, palaeogeography, and climate. The main factors that control the final composition of a sediment, besides the parent material defined by the source rocks, include chemical weathering (in situ as well as in transit, i.e., on the way from source to sink), erosion, mechanical alteration in transit and mixing of material derived from different sources. Sediment composition provides important information on potential source rocks as well as alteration in transit. Additional geological information on the provenance of ancient sediments comes from palaeogeography, facies patterns and transport directions. Data acquisition in provenance analysis can be grouped into three principal categories: (i) bulk composition, (ii) subcompositions, and (iii) single grains (varietal studies): (i) The bulk composition of sediments, both chemistry and framework petrography, is frequently used to determine the tectonic setting of sediment suites (e.g., Dickinson QFL-plots) or calculate the degree of chemical alteration (e.g., CIA, Chemical Index of Alteration). Information on grain size is crucial because sediment bulk composition is strongly related to grain size, and certain techniques are only applicable to certain grain-size fractions (e.g., petrographic point-counting). Petrographic techniques may provide the most detailed information because they account for both compositional and textural features of sediment grains. Chemical data, however, lack textural information but allow for a much faster and more precise determination of compositional variability (von Eynatten et al. 2003). (ii) The analysis of subcompositions allows concentration on information obtained from a specific group of sediment grains, for example, heavy minerals. Heavy minerals are high density constituents of sediments and usually need concentrating into mineral separates before quantitative analysis is performed. Heavy minerals are highly source-rock sensitive because they include many index minerals of, for example, metamorphic facies or grade. Similar to sediment bulk composition, heavy mineral assemblages in clastic sediments are modified during sediment erosion and dispersal, and later diagenesis. Hydrodynamically invariant heavy mineral ratios may provide excellent monitors of source area and climate dynamics (Morton and Hallsworth 1994). (iii) Varietal studies comprise an ever-increasing number of techniques related to the determination of lithology (e.g., tourmaline chemistry), metamorphic conditions (e.g., rutile geochemistry), formation ages (e.g., U-Pb dating) as well as cooling ages and rates (e.g., fission-track dating) of the source rocks. These techniques have the great advantage that they provide a direct link to the source rock without significant modification by alteration processes, i.e., the influence of factors other than source-rock characteristics being minimised. The most recent developments in this field are related to extracting multiple radiometric as well as chemical information from one and the same grain (e.g., zircon He-Pb double dating; Reiners et al. 2005). The selection of analytical techniques to solve provenance-related problems in basin analysis depends on the specific questions raised. In many cases, the most promising approach lies in a combination of single-grain varietal studies (to constrain source area lithology/petrology and geo-/thermochronology) and a bulk-composition technique (to gain further information on sediment mixing and alteration in transit). This combination is also best suited to relate the results of state-of-the-art provenance analysis to the bulk mass transfer from the source area to the sedimentary basin (Weltje and von Eynatten 2004).

rest on top of inverted structures. The “Subhercynian Unconformity” itself is slightly folded or is even the basal surface of secondary marginal troughs (e.g., Kockel 2003; Nielsen et al. 2005), so that the exact timing of inversion is still a matter of controversy (Grass and Geluk 1999; Voigt et al. 2004, 2006; Mazur et al. 2005; Nielsen et al. 2005; Krzywiec 2006b; von Eynatten et al. 2007). The basin centre remained relatively stable and is characterised by continuous sedimentation of chalk from Cenomanian to earliest Palaeocene. This is why the “Subhercynian unconformity” is not always interpreted to define the end of DC 8, but rather is seen as marking the significant change

from pelagic to terrigeneous deposits in the Palaeocene. DC 8 corresponds roughly with the “North Atlantic Cycle” of Jacquin and Graciansky (1998) which is bounded by late Berriasian and late Maastrichtian boundaries.

4.1.2.9 Depositional cycle 9: Tertiary (Palaeogene, Neogene) Depositional Cycle 9 starts with the Mid Palaeocene sea level rise (Ziegler 1990). The erosional unconformity at its base was caused by earlier Late Cretaceous inversion

Depositional history and sedimentary cycles in the Central European Basin System

of some basins (see Voigt et al. this volume) and possibly an additional eustatic sea level fall in latest Cretaceous and earliest Tertiary times. Production of chalky carbonates came to an abrupt end in Late Palaeocene times and gave way to clastic sedimentation in the rapidly subsiding North Sea Basin and adjacent areas. Eocene and Oligocene deposits spread over the inverted and upwarped southern margin of the CEBS. North Germany was situated at the southern margin of the deeply subsided North Sea Basin with thicknesses of Cenozoic strata up to 500-1000 m and 2000 m in the Hamburg area.

The 3–10 km thick sedimentary fill of the Central European Basin was derived from four different sources. These are in decreasing order of importance: 1. clastic deposits, eroded from bedrock of the surrounding highs and transported into the basin 2. carbonates, precipitated from marine water or by benthic and planktonic organisms 3. salt, anhydrite and gypsum, mainly evaporated from restricted marine water bodies 4. volcanics, mostly lava flows and tuffs (in the early stages of basin evolution)

There was a generally transgressive trend from Late Palaeocene (Thanetian) to Early Oligocene (Rupelian) times with deposition of shelf mudstones and shallow marine sandstone series. Broad coastal plains developed at the southern basin margin and gave way to the accumulation of thick lignite deposits. Sedimentation was frequently interrupted by sharp drops in sea level causing erosional hiatuses, especially at the bases of the Early, Middle and Late Eocene (Ypresian, Lutetian, Priabonian; Vail et al. 1977). The Early Oligocene Rupelian Shale extends far to the south and marks the Tertiary peak transgression.

Quantification of erosion and deposition as well as mass-balance calculations have to consider that only the clastic part of the sedimentary column originates exclusively from erosion of surrounding source areas, whereas the rest is provided by other sources, mainly from sea water. The relative proportions of these sources change depending on the position within the basin. While marginal sections contain predominantly clastic deposits, evaporites and limestones dominate in the basin centre and spread to the marginal depositional systems during periods of high sea level (e.g., Posamentier et al. 1988). For the CEBS, this was the case in the Late Permian, the Middle Triassic, the Late Jurassic and the Late Cretaceous. Clastic deposits prograded into the basin centres in the early continental stages of basin development (Early Permian, Early Triassic) and during episodes of relatively low global sea level (Late Triassic, Early Cretaceous, Cenozoic).

Thereafter, there was a strong regressive trend with sharp drops and rises in sea level and shifts in shorelines primarily caused by the waxing and waning of the ice sheets in the Antarctic and Arctic (Vail et al. 1977). From Miocene times, a depocentre developed in the southeastern North Sea area (Brückner-Röhling et al. 2005; Voigt et al. this volume). At the same time, the Rhenish and Bohemian Massifs were further uplifted. Parts of northeastern Germany, mainly in the Lower Elbe area, on the other hand, are still subsiding, thus indicating that the CEBS is an active basin (Ziegler 1990; Reicherter 2005; Sirocko et al. this volume).

4.1.3 Provenance of sediments in the Central European Basin In accordance with the stress field, which defines the position of subsiding and uplifting areas and the regional base-level, the Central European Basin was filled with sediments from different sources. Nevertheless, the cycles described above reflect not only tectonic evolution and changing sea level through time but also the mode of sedimentation depending on particular depositional environments. Additionally, climate and configuration of epeiric seas (open to the ocean or restricted) controlled which kind of sediment was deposited. Temporal and spatial distribution of deposits, sedimentation rates and sedimentary stacking patterns are the result of complex interactions of these driving forces (e.g., Coe 2003).

Clastic deposits such as sandstones and pelites were delivered from two main source areas with varying relative importance through time. Techniques to decipher the provenance of clastic sediments are summarised in box 4.1.2. The more or less continously uplifting Fennoscandian Massif in the north delivered large quantities of weathered bed rock during the entire basin history from the Late Carboniferous to the Cenozoic. The Fennoscandian Massif corresponds to the uplifted Proterozoic European Craton and extends to the southeast to the Russian Platform with a thin cover of Phanerozoic sediments. The Bohemian Massif in the southeast, which separated the Central European Basin from the Tethys was active as a source area from Permian to Recent with maximum erosion in the Permian and Early Mesozoic. This Massif forms the central region of the Variscan orogen and together with the Variscan basement of the Amorican Massif represents the southern source area of the Central European Basin System (e.g., Ziegler 1990). The influence of the intrabasinal Rhenish Massif on deposition is still under debate. In general its contribution to the total clastic sediment supply was very limited during Triassic and Jurassic times but became more important in Late Mesozoic to Cenozoic times (Murawski et al. 1983). Early Cretaceous deposits, for example, contain abundant coaly fragments that prob-

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ably derive from the erosion of Carboniferous coal-bearing strata in the Rhenish Massif (Littke et al. 1998).

The freshly exhumed basement rocks then form a new source of basin sediments.

During later stages of basin evolution, basin differentiation and inversion led to the formation of small-scale sub-basins such as the Lower Saxony Basin or the Subhercynian Basin as well as uplifted highs. In particular, uplifting basin parts acted during inversion as local source of sediments. However, sediments were also supplied by uplifted graben shoulders during rifting and push-up structures of strike slip faults (Ziegler 1990). Deposits of the local sub-basins mostly do not derive from sources outside the basin but consist of reworked older deposits of the basin system. Therefore, these deposits do not contribute to the overall sediment budget of the basin fill. If internal parts of the basin are eroded and the reworked sediment is deposited in the sub-basins, the term “basin cannibalism” is applied. In some cases, nevertheless, the basement of the Central European Basin was exhumed in areas of exceptionally strong uplift that may amount to more than 7 km in a few million years (e.g., Harz Mountains; Voigt et al. this volume).

The long lasting passage of Central Europe through low (and thus arid) latitudes (Fig. 4.4.1) and its position near the northern Tethys margin resulted in an exceptional thickness of evaporitic deposits compared to other basins (Bayer et al. this volume). Thick evaporites were deposited during the Early Permian (Upper Rotliegend), Late Permian (Zechstein), Triassic (Upper Buntsandstein, Middle Muschelkalk, Middle Keuper) and Late Jurassic, with Late Permian Zechstein evaporites being by far the thickest. The initial volume of these salt deposits is difficult to reconstruct because of significant salt redistribution (salt pillows and diapirs) and/or dissolution of extruded diapirs and exhumed salt deposits (Scheck et al. 2003b). Primary thickness of Zechstein evaporites may reach more than 1.5 to 2 km in some areas, suggesting that the basin was deep and underfilled before flooding. Decreasing subsidence rates and a lowering of the accommodation space allowed only relatively thin evaporite layers to accumulate in the Triassic and Late Jurassic. While the majority of Figure 4.1.5. General scheme of sediment generation, distribution and depositional style in a continent near depositional environment influenced by marine ingressions

Depositional history and sedimentary cycles in the Central European Basin System

salt deposits were precipitated from sea water, Permian spores occurring occasionally in Keuper salt and salt geochemistry suggest that salt domes pierced the surface and that some Permian salt was redeposited in Late Triassic continental settings (Trusheim 1960).

Each of the described sediment sources is characterised by various depositional systems, differing in the source of sediment, the mode of sediment input and distribution, and subsequent modifications during diagenesis (e.g., compaction, diapirism).

Carbonate deposits (limestones and dolomites) also contributed to the basin fill, particularly in Middle Triassic (Muschelkalk), Late Jurassic and Late Cretaceous. Carbonates originate either directly by precipitation from seawater (part of micritic limestones and ooliths) or from biogenic sources (shells and bacteria-induced precipitation). Of special note is the chalk, a biogenic porous carbonate composed of coccoliths, foraminifers and calcisphaeres which was deposited to enormous thickness during the Late Cretaceous (Ziljstra 1995). Chalk is restricted to pelagic environments far from source areas (Hancock 1975). Total thickness of carbonates may add up to more than 1000 m in the basin centre and a few thousand meters in some sub-basins, for example the eastern Lower Saxony Basin or the Danish Basin.

Figure 4.1.5 shows how sediment generation, distribution and stacking patterns (“depositional style”) of such depositional systems may vary, using four important examples: 1. Terrigeneous depositional systems, both continental and marine, show sediment input from the basin margins (alluvial fans or deltas). The sediment source is located outside the depositional system. The grain size of deposits decreases in a basinward direction, accompanied by a gradual transition from conglomerates to sandstones and pelitic rocks towards the basin centre. Lateral shift of river channels or longshore transport by marine currents produce facies belts running parallel to the basin margin. 2. Carbonate ramps, the main type of carbonate depositional systems in the Central European Basin, are characterised by a high-energy facies belt (banks and shoals of oolithic, peloidal or skeletal sands), where a major part of the sediment is produced. Minor sources are terrigeneous input (clay) and micritic carbonates that precipitated from sea water (especially lagoonal areas and tidal flats). Wind-driven currents and storms disperse sediments both land- and basinward. Stacking patterns are controlled by the migration of the narrow grainstone belt. It is surrounded by extended fine-grained carbonate mud deposited in lagoons and on tidal flats landwards and marls with storm-beds towards the basin centre (e.g., Aigner 1985). 3. A completely different depositional style is observed in pelagic depositional systems (chalk). The whole sediment consists of the skeletal remains of planktonic organisms (coccoliths, foraminifers, calcisphaeres) settling down from the water-column. The pure carbonate is produced within the system and as planktonic rain drapes both subsiding areas and swells similarly. Indistinct facies belts and gradual thickness changes can be observed due to higher productivity in nutrient-rich water. Redistribution of sediments occurs only if currents (mostly storm-induced) reach the sea bottom in periods of lower sea level or if the steepness of slopes allows gravitational mass flows. Aggradational stacking patterns dominate the entire system (e.g., Hancock 1989). 4. Evaporitic depositional systems reflect changes in brine chemistry within the basin due to evaporation and precipitation of components according to decreasing solubility. Carbonate deposits (carbonate platforms or ramps) of the initial stages of evaporation are rimmed by prograding sulphate ridges. Subsequently,

Towards the basin margins, limestones often pass into a mixture of fine-grained clastics with carbonates (marls and calcareous siltstones) and may reach considerable thickness, especially in Jurassic to Cretaceous time. Taking into consideration that carbonate content of marls is mostly higher than 50%, clastic rocks may also contain high amounts of carbonate derived from sea water. Intensive volcanism accompanied the initial phase of basin evolution and left behind thick volcanic rocks and pyroclastic deposits of Early Permian age. These volcanics reach an enormous thickness in the eastern-central part of the basin (Breitkreuz et al. this volume). Traces of active volcanic sources were also found in the Late Cretaceous and in the Cenozoic, but such relatively thin tuffs do not contribute significantly to the basin fill. Throughout the evolution of the CEBS, episodes of predominantly clastic deposition alternate with phases of evaporitic and biogenic deposition. Sedimentation types were governed by regional tectonics, climate and sea level. Terrigeneous input was strongest in the early stages of basin evolution (Permian) as well as in the Early Triassic, whereas limestones deposition prevailed in the Middle Triassic, Late Jurassic and especially in the Late Cretaceous, when sea level reached the highest level in earth history. Thick evaporites were precipitated during arid climate periods when the basin was occupied by restricted marginal seas. The highest sedimentation rates are observed during phases of evaporite deposition (300 m/Ma) and not when clastic deposition was predominant (20–100 m/Ma), except for strongly subsiding sub-basins (see box 4.1.1).

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the remaining space is filled with halite and (to a minor degree) K- and Mg-salts. Thick evaporitic deposits require a vast accommodation space (deep water) and brine exchange with the open ocean (barred basin). Evaporitic facies belts show either a “bull’s eye” pattern if the basin was closed or a “tear drop” facies pattern in the case of a semiclosed evaporite basin (e.g., Kendall and Harwood 1996). One important type of depositional system, namely carbonate platforms, characterised by flat and shallow lagoons fringed by steep reefs and coarse fore-reef breccias is not present in the sedimentary history of the CEBS. Different architecture of sediment bodies and fluctuating sources have to be taken into account when the sub-

sidence- and tectonic history of a basin is investigated. Misinterpretations of seismic sections or structural situations may arise from simple interpretations of on- and offlap patterns as progradational sequences. These patterns might also represent thinning on top of swells or secondary effects due to salt movement (pillows, diapirs, piercing and solution). Deposition of chalk may even occur on top of rapidly uplifting swells due to high sea level and thus may mask the timing of tectonic events. Evaporites appear as nearly chronostratigraphic units in a basin whereas clastic sediments (especially if coarsegrained) in general need a lot of time to prograde. Both spatial and temporal patterns of these deposits differ significantly and could lead to false conclusions regarding basin evolution. Some of these aspects will be highlighted in the next chapters.

4

Chapter 4.2

Basin initiation: Volcanism and sedimentation Ch. Breitkreuz · M. Geißler · J. Schneider · H. Kiersnowski

4.2.1 Late Palaeozoic basins in central Europe – distribution, volcanic activity and magmagenetic aspects During the Late Palaeozoic, the Central European Basin System (CEBS) was subdivided into the Northern Permian Basin (NPB) and the Southern Permian Basin (SPB, Fig. 4.2.1). In its initial phase, at the Carboniferous-Permian transition, the NPB and in particular the SPB experienced strong volcanism. The CEBS was one of many zones in Europe with Late Palaeozoic volcanic activity (Schneider et al. 1995c; Cortesogno et al. 1998; von Raumer 1998; Deroin and Bonin 2003; Timmerman in press). Connected to the NPB, the north-south striking calc-alkaline to alkaline Oslo Graben system was active (Neumann et al. 1992, 2004). To the south of the SPB, a number of intermontane basins existed, that were related to transtensional tectonics in the Variscan orogen and which developed in different ways. For example, in the Saar-Nahe Basin (Lorenz and Haneke 2004; Seckendorff et al. 2004), the Saale Basin (Ehling and Breitkreuz 2004) in Germany and in the Intra-Sudetic and Northern Sudetic Basin in Poland (Awdankiewicz 1999, 2004) volcanic activity occurred only late in the basin evolution (Fig. 4.2.2). In contrast, opening and fill of the pull-apart type Collio Basin in the Italian Alps was accompanied by mainly explosive volcanism (Breitkreuz et al. 2001). Similarly, the early history of the Thuringian Forest Basin in Germany was accompanied by extensive magmatism (Andreas et al. 2005). Accordingly, the Late Palaeozoic extensional and transtensional basin systems of Europe can be grouped into basins with initial magmatism and those with late or no volcanism. The former type is similar, for example, to the evolution of the Basin-and-Range Province in the western United States (Gans et al. 1989; Christiansen and McCurry 2008) for which decompression of mantle lithosphere, “fertilised” during previous subduction, had been assumed as a model (Harry and Leeman 1995; Hawkesworth et al. 1995). For the Saale Basin, the presence of mantle magma mixed with continental crust with a subduction fingerprint has been demonstrated (Romer et al. 2001; for Saar-Nahe Basin see Schmidberger and Hegner 1999).

The discussion of CEBS magma generation involves models of intracontinental wrench tectonics and postorogenic extension (Marx et al. 1995; Benek et al. 1996), crustal delamination (Schott and Schmeling 1998) and mantle plume activity (Neumann et al. 2004; Torsvik et al. 2008). During the Upper Rotliegend II (Parchim Formation), small basaltic volcanic fields developed in various regions of the SPB (review in Marx et al. 1995).

4.2.2 Data base, distribution and volumes of Late Palaeozoic volcanics in the CEBS With the exception of outcrops in the FlechtingenRoßlau Block near Magdeburg, information on volcanic deposits in the NPB and the SPB comes from hundreds of hydrocarbon exploration wells up to 8 km depth and from the interpretation of seismic data. In the area northwest of and around Hamburg Late Palaeozoic rocks were deeply buried beyond the reach of hydrocarbon exploration wells (Fig. 4.2.1). The data set is strongly biased by well distribution and penetration depth. Furthermore, the thick Zechstein evaporates veil Rotliegend structures in many regions hampering clear seismic imaging (Krawczyk et al. 1999). Nevertheless, abrupt changes in thickness and facies testify to active extensional tectonics throughout much of the latest Carboniferous to Late Permian. Repeated progradation of coarse clastic sedimentary systems have been interpreted as tectonic pulses (Gebhardt et al. 1991). CEBS volcanic rocks display calc-alkaline to alkaline composition of intra-plate affinity (Marx et al. 1995, Benek et al. 1996; Maliszewska et al. 2003). Due to burial and hydrothermal overprint the rocks experienced strong alteration in places (Brecht 1999; Wolfgramm and Schmidt Mumm 2001; Maliszewska et al. 2003). Volcanogenic jointing and cleavage during emplacement and cooling as well as Late Palaeozoic to Cenozoic tectonic overprint (Hecht et al. 2003) led to abundant fracture porosity. In the NPB, volcanism was dominated by basaltic successions and associated subvolcanic bodies (Vejbæk 1990;

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Figure 4.2.1. Distribution of Rotliegend volcanic and sedimentary rocks in the Central European Basin System (CEBS) (modified after Heeremans et al. 2004 and others); wells: Am 1/68 – Angermünde 1/68, Dgf Z1a – Dageförde Z1a, Fdl 1/71 – Friedland 1/71, Gap 1/86 – Garlipp 1/86, Grt 2/75 – Gartz 2/75, Gür 3/76 – Grüneberg 3/76, Kotz 4/74 – Kotzen 4/74, MirNs 1/74 – Mirow 1/74, Ob 1/68 – Oranienburg 1/68, Obje IG1 – Objezierze IG 1, Pa 1/68 – Parchim 1/68, Pap 28 – Paproc´ 28, Pkn 1/71 – Penkun 1/71, Pud 1/86 – Pudagla 1/86, Sok 1 ´ ˛ 1; volcanic sub-provinces: MVSP = Mecklenburg– Santok 1, Tuc 2/76 –Tuchen 2/76, StrKr IG1 – Strzelce Krajenskie IG 1, Zien 1– Zielecin Vorpommern Sub-province, EBSP = E Brandenburg Sub-province, FASP = Flechtingen-Altmark Sub-province; WBH = West Brandenburg High; for position see figure 4.2.2

Heeremans and Faleide 2004). For the SPB, in the North Sea area and in the area of northwestern Germany including the border area with the Netherlands, the distribution of volcanic rocks is mainly limited to NNW-trending Upper Rotliegend II grabens (Fig. 4.2.1), and their thickness varies from a few meters to a few hundred meters. The volcanic rocks are dominated by basaltic and andesitic lava flows, with a minor presence of rhyolitic lavas and pyroclastic deposits (Marx 1995; Gast and Gundlach 2006; Börmann et al. 2006). The thickest volcanic successions are known from the area within the central SPB north of Berlin with a thicknesses up to 2250 m (e.g., well Friedland 1/71 stopped after 2249 m of volcanic rocks without reaching the base

of the volcanic pile). Benek et al. (1996) calculated a minimum volume of about 50,000 km³ for the area of the central SPB. Considering the entire SPB, the total volume of volcanics is probably on the scale of 80,000 km³. For comparison, the Cenozoic San Juan Volcanic Field in the western United States has a volume of more than 40,000 km³ (Lipman et al. 1978). In western Poland the distribution of volcanic rocks extends towards the area of Poznan´ (Fig. 4.2.1). Volcanic facies includes ignimbrite sheets, lava domes and associated pyroclastic deposits as well as andesitic lava. Numerous hydrocarbon wells document a thinning of the volcanic succession towards the east (summaries in Jackowicz 1994; Maliszewska et al. 2003).

Basin initiation: Volcanism and sedimentation

Figure 4.2.2. Palaeogeographic position of important Permo-Carboniferous basins in central and southern Europe: AU – Autun Basin, AVD – Athesian Volcanic District, BLG – Blanice Graben, BCG – Boskovice Graben, BU – Bourbon l’Archambault Basin, CA – Carpatian Basin, CB – Collio Basin, CR – Carnic Alps, DB – Donezk Basin, DÖ – Döhlen Basin, EB – Erzgebirge Basin, FL – Flechtingen-Roßlau Block, FR – Franconian Basin, GP – Guardia Pisano Basin, IF – Ilfeld Basin, IS – Intra-Sudetic Basin, KP – Krkonoše Piedmont Basin, LC – Lu Caparoni Basin, LO – Lodève Basin, MO – Montceau les Mines Basin, NS – North Sudetic Basin, PBF – Pays de Bray Fracture, PD – Perdasdefogu Basin, RGL – Rhein Graben Lineament, RÜ – Rügen, SB – Saale Basin, SNB – Saar-Nahe Basin, SPBV – Southern Permian Basin Volcanic Zone, ST – St. Etienne Basin, SV – Salvan-Dorénaz Basin, TF – Thuringian Forest Basin, WCB – Western and Central Bohemian Basins, WEI – Weissig Basin, ZÖ – Zöbingen; Tectonic lineaments: EBL – Elbe Lineament, FRL – Franconian Lineament, GSH – Grand Sillon Houllier Fracture Zone, HRF – Hunsrück Fracture Zone, TTZ – Tornquist-Teisseyre Zone, VDF – Variscan Deformation Front (modified after Roscher and Schneider 2006)

4.2.3 Stratigraphy and geochronology of volcanic successions in the SPB The strong volcanic activity prior to the main sedimentary development of the SPB can be subdivided into two peri-

ods. During the first main period (latest Carboniferous to Lower Rotliegend) extensive magmatism with basaltic/ andesitic to rhyolitic geochemical composition occurred along major tectonic faults, graben- and half-graben structures, probably related to NW-SE trending strike-slip and wrench tectonics, and subordinately probably along volcanic ring structures. Prior to the main volcanic phase,

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during the Stephanian, sedimentation and minor explosive volcanism is known from small depressions of Lower Rotliegend time, as in the Süplingen area of the Flechtingen-Roßlau Block. Some wells exposed conglomerates which underlie the volcanic succession and comprise exclusively sedimentary and metamorphic clasts of the underlying Variscan basement. Later, in the course of the Rotliegend (Upper Rotliegend II), a second period of magmatism created small isolated volcanic fields of basaltic geochemical composition. The magma rose mainly along NNE-SSW faults, related to the Altmark tectonic movements (Bachmann and Hoffmann 1995). Local volcanic centres are known from the Horn Graben in the North Sea area (Marx et al. 1995; Glennie et al. 2003), the eastern Netherlands (Verdier 1996), the Soltau area between Hamburg and Hannover (Marx 1995), and from the Garlipp and Kotzen areas north of the Flechtingen-Roßlau Block (Geißler et al. 2008; Fig. 4.2.1). Numerous attempts of K/Ar dating of SPB volcanic rocks led to ambiguous, rather young ages (many unpublished data), presumably the result of thermal overprint during the Triassic-Jurassic (Zwingmann et al. 1998; Brecht 1999; Jacobs and Breitkreuz 2003). Breitkreuz and Kennedy (1999) and Breitkreuz et al. (2007) provided reliable emplacement ages between 290 and 302 Ma with a focus between 295 and 299 Ma. These ages were obtained by SHRIMP dating of zircons separated from silica-rich volcanic and sub-volcanic rocks drilled in Denmark, Germany and Poland. Thus, the climax of volcanic activity in the SPB took place at the Carboniferous-Permian transition.

4.2.4 Volcanic facies in the SPB A wealth of petrographic and geochemical data on the volcanic rocks has been summarised in numerous unpublished reports of the hydrocarbon industry. Combined with new observations, a portion of this dataset has been published (Benek et al. 1973, 1992, 1996; Korich 1992; Hoth et al. 1993a; Jackowicz 1994; Marx et al. 1995; Maliszewska et al. 2003). Volcanic facies analysis on drill cores, the advantages and limitations of which are featured in Box 4.2.1 and figure 4.2.3, is the focus of this contribution. The whole SPB volcanic area comprises basaltic lava fields, Mg-andesite shield volcanoes, extended ignimbrite sheets and other pyroclastic deposits, as well as silica-rich lava domes (both as isolated small bodies and as voluminous complexes). In addition, subvolcanic andesitic complexes (Awdankiewicz et al. 2004) and silicarich laccoliths or single intrusions have been recognised (Paulick and Breitkreuz 2005; Geißler et al. 2006). For the entire SPB, Geißler et al. (2008) listed the following main types of volcanic products, from abundant to subordinate (see also Fig. 4.2.4): a. b. c. d.

SiO2-rich lava flows, lava domes and laccoliths SiO2-rich ignimbrite sheets (welded and non-welded) Basaltic to andesitic shield volcano complexes Pyroclastic deposits related to the eruption of lava domes (fall-out tuffs, tuff breccias, and minor surge and block-and-ash flow deposits) e. Smaller basaltic and andesitic lava fields f. Basaltic to andesitic sill complexes g. Volcanoclastic sediments

Box 4.2.1 Working with fossil volcanic successions exposed in drill cores One advantage of working with drill cores in fossil volcanic zones is the chance to carefully observe the top contact of coherent volcanic bodies. Of stratigraphic importance is the distinction between extrusive and intrusive facies. This has been elaborated for andesitic units in the Flechtingen-Roßlau Block (Fig. 4.2.1) by Awdankiewicz et al. (2004). With intrusive contacts the host sediment textures provide information on the age relation between intrusion and host (soft vs. lithified host; see also Kokelaar 1982; Orth and McPhie 2003). In this regard, Geißler et al. (2006) published preliminary results for the central SPB, Mock et al. (2005) for the eastern Saale Basin. Thick SiO2-rich lava, laccolith or ignimbrite units typically display a strong internal textural variation which developed during emplacement, cooling and diagenetic alteration (Allen 1988; McPhie et al. 1993; Gifkins et al. 2005). Especially with incompletely cored successions, delineating emplacement units in one well and correlation of units between neighboring wells is a challenge. Many SiO2-rich emplacement units like lava domes and laccoliths have high aspect ratios (height vs. length) implying abrupt lateral changes in thickness and facies variation. For the Late Palaeozoic volcanic rocks in the CEBS, a combination of detailed volcanic facies analysis, evaluation of geophysical logs and of immobile trace element ratios (Fig. 4.2.3) proved to be useful in many cases (see also Paulick and Breitkreuz 2005). Underlying, interfingering and overlying sediments should be studied in close context with the volcanic units for the following reasons: (i) Intrusive bodies typically are under- and overlain by sediments of similar facies, whereas the emplacement of lava and ignimbrite sheets often influences the sedimentary environment (e.g., Buesch 1991). (ii) Volcanoclast content in the sediments can provide information on volcanism which had been active beyond the drill site. (iii) The sedimentary environment may provide an indication of climatic and hydrological conditions during the emplacement of volcanic units and the geotectonic position. (iv) Sediments incorporated at the base of extrusions or surrounding intrusions may affect the geochemical composition of the carapace facies part of volcanic bodies.

Basin initiation: Volcanism and sedimentation

Figure 4.2.3. Work steps of lithological interpretation of drilled volcanic rocks, exemplified by the well Penkun 1/71 (for location see Fig. 4.2.1; Hoth et al. 1993b, Paulick and Breitkreuz 2005); a) lithological classification based on drill cores and thin sections; photos between a) and b) show some typical volcanic facies types (width of photographs is 6 cm); b) gamma ray log, green arrows indicate apparent lithological changes; c) the ratio of the immobile elements Ti and Zr helps to delineate lithological units; d) final lithological interpretation: the succession is dominated by two rhyodacitic dome complexes, the lower one being intruded by an andesitic sill (U/Pb-ages taken from Breitkreuz and Kennedy 1999 and Breitkreuz et al. 2007)

Considering semiarid to arid climatic conditions within the SPB during the Permian time and the knowledge of weathering resistance of different volcanic edifices, a, c and f are presumed to have created substantial topography, whereas b, d, e and g largely levelled pre-existing relief. Accordingly, in NE Germany, the area of the greatest known thicknesses of volcanics, three sub-provinces can be outlined: these are the Mecklenburg-Vorpommern Sub-Province (MVSP), the E Brandenburg Sub-Prov-

ince (EBSP), and the Flechtingen-Altmark Sub-Province (FASP; Fig. 4.2.1). The MVSP and the EBSP are dominated by topography-forming volcanism, whereas much of the Altmark area and the Flechtingen-Roßlau Block are dominated by topography-levelling ignimbrite sheets (Geißler et al. 2008; see chapter 4.2.6). Compared to many Neogene intracontinental volcanic fields, the SPB volcanic zone is characterised by the predominance of SiO2-rich lava domes and subvolcanic intrusions over ignimbrites.

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Ch. Breitkreuz · M. Geißler · J. Schneider · H. Kiersnowski Figure 4.2.4. Selected volcanic lithofacies types (rock slabs) of the Southern Permian Basin (further examples were displayed by Paulick and Breitkreuz 2005 and by Geißler et al. 2008); a) core facies of a porphyritic SiO2rich lava, well Strzelce Krajenskie ´ IG 1 at 4204 m; b) flow foliated SiO2-rich lava with small phenocrysts, well Dageförde Z1a at 4873 m; c) strongly porphyritic SiO2-rich lava with flow foliation, well Garlipp 1/86 at 4566.2 m; d) spherulitic upper carapace facies of SiO2-rich lava(dome), well Kotzen 4/74 at 5437.6 m; e) lithic-bearing ignimbrite, well Paproc´ 28 at 2811 m; f) crystal- and lithic-rich tuff, well Kotzen 4/74 at 5390.5 m; g) laminated fine tuff with few small lapilli, well Gartz 2/75 at 3973 m; h) vesicular Mg-andesite lava, well Santok 1 at 3491.5 m; i) tectonic breccia in core facies of a Mg-andesite lava flow, well Oranienburg 1/68 at 4462.9 m; j) irregular top breccia of a Mg-andesite lava flow, well Oranienburg 1/68 at 4362.7 m; k) calc-alkaline basalt lava: the large vesicle shows complex diagenetic/ hydrothermal filling, well Garlipp 1/86 at 4406 m; l) vesicular tholeiitic basalt lava (strongly oxidised by diagenetic/hydrothermal overprint), well Kotzen 4/74 at 5160.5 m; width of each photograph = 5 cm

4.2.5 Syn- to postvolcanic sedimentation during the Lower Rotliegend and Upper Rotliegend I Due to spacial isolation of outcrops and wells and the scattered biostratigraphic data and facies analysis, stratigraphy remains poorly defined for the early SPB sediments. The sedimentary history within the SPB started in the latest Carboniferous with the re-sedimentation of Early Carboniferous rocks (e.g., magmatic and metamorphic rocks, greywackes, mudstones) accumulating mainly as coarse clastic deposits in local small basins (Type I, Table 4.2.1).

Presumably, onset of volcanic activity was diachronous in different regions of the SPB. Thus it is conceivable that Type I and II sedimentation (Table 4.2.1) took place at the same time in the early SPB evolution. Synvolcanic Type II-sediments are preserved in a few areas only. Examples of this sedimentation are to be found in wells within the Barnim Basin northeast of Berlin, in outcrops of the Flechtingen-Roßlau Block and in a few wells in the Altmark area as well as in the Strelasund Depression south of Rügen (Fig. 4.2.1). Characteristic of these deposits is a remarkably high amount of lacustrine and/or carbonate sediments, and the intercalation of pyroclastic horizons (Gaitzsch 1995; Schneider and Gebhardt 1993; Gaitzsch et al. 1998).

Basin initiation: Volcanism and sedimentation

Table 4.2.1. Upper Carboniferous to Latest Rotliegend I sediments of the SPB

·

Lithology

Spacial relation to the volcanic complexes of the Altmark Subgroup

Temporal relation to volcanic activity: example wells / outcrops

Type I

Fluvial, coarse clastic, wellrounded conglomerates; clasts resemble consolidated Carboniferous sediments

Below the volcanic succession

Prevolcanic, pre-Altmark Subgroup: Northeast of Berlin: Angermünde 1/68 and Oranienburg 1/68 (Hoth et al. 1993b; Geißler et al. 2008) Strelasund Depression: Pudagla 1/86 (Hoth et al. 1993b) Flechtingen-Roßlau Block: Basal Süplingen Formation (Gaitzsch et al. 1995)

Type II

Lacustrine, alluvial to fluvial, fine to coarse clastic, mainly carbonatic, varying amounts of volcanic detritus and layers (tuff, pumice fragments)

Within basins in more or less close vicinity of active volcanoes

Synvolcanic, Altmark Subgroup: Barnim Basin: Grüneberg 3/76, Tuchen 2/76 (Gaitzsch 1995) Flechtingen-Roßlau Block: Upper Süplingen Formation, Eiche- and Bodendorf members, Bebertal Formation (Gaitzsch et al. 1995) Northwest of Poznan: ´ Objezierze IG 1 (Maliszewska et al. 2003)

Fluvial-lacustrine, coarse to fine clastic and partly carbonatic, with late eolian influence in places

Above or between extinct volcanoes

Type III

· · · · · ·

In the ensuing time of major volcanic activity, sedimentation was reduced to depressions between volcanic edifices and depressions resulting from Lower Rotliegend tectonics. Later, in the course of the Lower Rotliegend and Upper Rotliegend I, filling of existing volcanic and tectonic depressions continued. The ongoing stepwise covering of volcanic edifices and tectonic highs culminated in the SPB main sedimentary evolution in association with intensified regional subsidence starting in Upper Rotliegend II time (see next chapter). In some wells a continuous succession of syn- to postvolcanic sediments (Type II to III, Table 4.2.1) is exposed which displays no significant change in depositional style (e.g., Grüneberg 3/76, Table 4.2.1). We define the Upper Rotliegend I deposits (Müritz Subgroup, Type III) as coarse to mainly fine clastic sediments (see also Kleditzsch 2004) similar to the Lower Rotliegend sediments (that is, a relatively high content of carbonate within the clastic sediments). In contrast to the findings of previous studies (e.g., Schneider et al. 1995b), we observed no pyroclastic material in the Upper Rotliegend I deposits. In the upper part, playa-like deposits are predominant, and they transit into sandy deposits with signs of eolian input. In places, the Upper Rotliegend I succession is followed by the onset of coarse clastic alluvial fan deposition related to the Parchim Formation.

· ·

Postvolcanic, Müritz Subgroup: Mirow 1/74, Parchim 1/68 (Geißler et al. 2006) ˛ 1 (Maliszewska et al. 2003) Zielecin

4.2.6 Landscape evolution during the initial phase of the SPB The SPB evolved on top of the Variscan thrust and fold belt and its northern foreland (Krawczyk et al. 1999; McCann 1999). Thus during the initiation of the SPB, in the area to the south of the Variscan Deformation Front (VDF, Fig. 4.2.2) the presence of a differentiated topography can be assumed. To the north of the VDF a rather flat fluvial to lacustrine environment may be envisioned. It was only at the northern margin of the SPB that tectonic activity along the Tornquist-Teisseyre Zone (TTZ, Fig. 4.2.2) led to differential subsidence in narrow WNW-ESE trending depressions (Rieke et al. 2001). Related to intense dextral strike-slip movements (Franconian tectonic movements, Bachmann et al. this volume), that affected large areas of Palaeo-Europe at the Carboniferous-Permian transition (Arthaud and Matte 1977; Ziegler 1990), the area of the central SPB was subject to intense volcanism. The Altmark Subgroup volcanism and associated tectonic activity changed the above-described landscape dramatically within a few million years. In some marginal regions of the SPB, such as the North Sea, Lower Saxony and western Poland, tectonic control of landscape evolution

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dominated over volcanic effects. However, in the central SPB (Fig. 4.2.1), intense volcanism led to the formation of i) an ignimbrite-dominated province in the Altmark region (FASP), ii) a shield volcano complex in the Berlin area (EBSP, see also Benek et al. 1996) and iii) an area dominated by voluminous lava domes in eastern Mecklenburg-Vorpommern (MVSP, see also Paulick and Breitkreuz 2005). In the latter two regions weathering-resistant, topography-forming edifices developed, resulting in substantial topography (Geißler et al. 2008). Intra-basinal structural highs, like the West Brandenburg High and the NE Brandenburg - Wolsztyn

High (Fig. 4.2.1), caused variations in thickness and absence of volcanic rocks in places. In the course of the Rotliegend, the volcanic-topography was partially eroded and incrementally covered with sediments. In the border region between northeastern Germany and northwestern Poland, covering was not complete until the Zechstein (Roch et al. 2005). Geißler et al. (2008) emphasised that the prominent Rotliegend hiatus present in the central SPB (Bachmann et al. this volume; Menning et al. 1988; Hoffmann et al. 1997) can be ascribed in parts to these volcano-topographic effects.

4

Chapter 4.3

Upper Rotliegend to Early Cretaceous basin development H. Stollhofen · G.H. Bachmann · J. Barnasch · U. Bayer · G. Beutler · M. Franz M. Kästner · B. Legler · J. Mutterlose · D. Radies *

4.3.1 Introduction The term Central European Basin System (CEBS) addresses a broad, WNW-ESE trending depositional area extending from Central England to SE Poland (Maystrenko et al. this volume) as well as the North Sea from offshore Norway to the Netherlands. It is composed of a series of structurally controlled sub-basins (Scheck-Wenderoth and Lamarche 2005) that acted as local depocentres. The heterogeneity in distribution and timing of deformation in the CEBS was probably controlled by pre-existing faults and other factors (e.g., variations in salt thicknesses) which caused local stress perturbations (Lohr et al. 2007a). The CEBS area cannot be viewed as a single basin in the strict sense, but as a combination of tectonically different types of basins with varying outlines, partly stacked one on top of the other, developing a cumulative geometry, which is commonly referred to as the CEBS (cf. Brink et al. 1992; Maystrenko et al. 2005b). Some of the evolutionary steps of the CEBS compare well with the development of African Karoorifts involving sequential stages of extensional faulting, heating and fault-controlled subsidence succeeded by subsidence related to periodic thermal cooling and contraction. A comparative example is the North Sea Basin, recognised by Sclater and Christie (1980) from backstripped subsidence curves, involving two superimposed phases of rifting and thermal cooling on top of one another. The present North Sea “basin” geometry is superficial. The CEBS is overlies the NE-SW trending Foreland Basin of the Variscan Mountain Belt which developed during Late Carboniferous times. Lower Rotliegend to Upper Rotliegend I (Early Permian) development then relates to a separate phase of transtensive pull-apart basin formation associated with a period of short-lived but intense bimodal volcanism followed by thermal subsidence (Breitkreuz et al. this volume; Bachmann and Hoffmann 1997; van Wees et al. 2000).

This chapter will review the Upper Rotliegend II (late Early Permian) to Early Cretaceous history of tectonics and sedimentation in the area of the CEBS, with a special focus on the North German Basin. The succession is framed by the widespread pre-Upper Rotliegend II unconformity at its base and terminates with the onset of Late Cretaceous/Early Tertiary inversion (Voigt et al. this volume) at its top. It subdivides into four further phases of extension and basin development referred to as (i) the Altmark “tectonic pulses” (Upper Rotliegend II), (ii) the Hardegsen phase (Middle Buntsandstein), (iii) the Early Cimmerian phase (Middle Keuper) and (iv) the Middle to Late Cimmerian phases (Middle Jurassic/Early Cretaceous). All of these extensional phases are associated with unconformities and may coincide with major depositional cycles (DC 3-8) (Bachmann et al. this volume; Fig. 4.1.3). Although many of the unconformities reflect basinscale tectonic, thermal and isostatic processes, they may display great local complexity and great variability on a regional scale. Following Kyrkjebø et al. (2004) for instance, the so-called “Late Cimmerian” unconformity is recorded by a complex set of angular unconformities, disconformities and nonconformities in space and time. The term “Cimmerian” was originally introduced by Suess (1909), based on a tectonic phase in Crimea at the Jurassic-Cretaceous boundary. This view was later refuted by Arkell (1956), and the transfer of tectonic phase from an active system to Central Europe has repeatedly been criticised. Nevertheless, the term survived in the literature (e.g., Ziegler 1990) and up to six “Cimmerian phases” have now been recognised lasting from the late Triassic to the Jurassic/Cretaceous boundary. Many of the associated unconformities, however, are only of regional to local extent. Therefore preference should be given to more neutral terms like (Early) Late Triassic, middle Aalenian, and Oxfordian-Ryazanian unconformities without genetic interpretation. Even the widespread “base Cretaceous unconformity” may consist of a set of merging regional unconformities within the CEBS, ranging from Oxfordian to early Cretaceous age (Kyrkjebø et al. 2004).

*with contributions by R. Littke · Y. Maystrenko · T. Voigt · J. Winsemann

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4.3.2 Upper Rotliegend II Stratigraphic framework The basal Upper Rotliegend II stratigraphic framework is pinpointed by the identification of the Illawarra magnetic reversal (Fig. 4.3.1) in the lowermost Parchim Formation (Havel Subgroup) in northern Germany (Menning et al. 1988, 2006). The Illawarra reversal has been assigned to the Middle to Late Wordian (Steiner 2006) and latest Wordian (Glenister et al. 1999), respectively. Based on magnetostratigraphic investigations, Steiner (2006) calculated an age of ∼267 Ma for the Illawarra reversal. Dating of a tuff located close to the Illawarra reversal in the Middle Permian (Guadalupian) reference section in West Texas/US (Glenister et al. 1999; Menning 2001) revealed a U/Pb ID-TIMS age of 265.3±0.2 Ma (Bowring et al. 1998). Consequently, there is a large time-stratigraphic gap between the Müritz Subgroup (Upper Rotliegend I) and the base of the Havel Subgroup (Upper Rotliegend II) lasting

as long as 15 Ma, from ~282 Ma to ~267 Ma (Menning et al. 2006). The Upper Rotliegend II comprises the Havel and Elbe Subgroups; equivalents of the latter are also referred to as “Weissliegend” (e.g., Glennie and Buller 1983). Palaeogeography and depositional environments Upper Rotliegend II strata of the Southern Permian Basin (SPB) extend E-W from Poland to west England and N-S from the German/Danish border to northeast of Frankfurt/Germany (Fig. 4.3.2). The basin centre was located in northern Germany, accommodating a succession of up to 2000 m thickness. A stepwise enlargement of the depositional area from this centre toward the west and the east can be observed over time. The Upper Rotliegend II succession is characterised by continental siliciclastics and minor evaporites which were deposited under arid to semi-arid climates (Glennie 1972, 1983). Four facies associations dominate the succession which have been in-

Figure 4.3.1. Lithostratigraphy of the German Upper Rotliegend II (modified from Plein 1995) and control parameters on deposition. Note, that here the base of the Wustrow Member is taken as the boundary between Dethlingen and Hannover Formations. The timing of the Elbe Subgroup is based on Gast (1995). Position of Illawarra reversal is according to Steiner (2006)

Upper Rotliegend to Early Cretaceous basin development

terpreted to reflect deposition in ephemeral fluvial (wadi), aeolian, sabkha and lacustrine environments (Glennie 1990; George and Berry 1993; Strömbäck and Howell 2002; Legler 2006). Particularly the wadi systems at the southern basin margin channelled sediment influx (Plein 1978; Gast 1988), but, following Gast et al. (1998) and Rieke (2001), an additional significant input from the east is also obvious. Aeolian deflation and re-deposition by prevailing easterly winds played an important role in sediment distribution. These aeolian deposits form the major plays targeted by Rotliegend gas exploration. They accumulated on a belt of saline mudflats and sandflats, fringing a perennial saline lake that existed in the basin centre since the Dethlingen Formation (Gast 1991). A characteristic feature of Upper Rotliegend II deposition is evidence of short-termed marine ingressions (Legler et al. 2005). The basin did not become fully marine, however, until the onset of the major Zechstein transgression.

Geodynamics and basin tectonics With the onset of the Upper Rotliegend II the longlasting phase of thermal subsidence that had followed Lower Rotliegend bimodal magmatism during the Upper Rotliegend I (Bachmann and Hoffmann 1997; van Wees et al. 2000) became superposed by a new phase of extensional faulting. In central and northern Germany the Upper Rotliegend II strata sit above a widespread erosional unconformity (cf. Schneider et al. 1995a; Gebhardt and Plein 1995) which was most probably induced by a re-orientation of plate boundary forces (cf. Schneider et al. 1995a; Stollhofen 2007). It is important for the understanding of the geodynamic context to note that the 267-265 Ma age of the Inge Volcanics of the Central Graben (Ineson 1993; Cameron 1993) seems roughly to coincide with the onset of Upper Rotliegend II deposition just prior to the Illawarra reversal at ~267 Ma (Steiner 2006). Similarly, in NE Germany at least two of the four Altmark tectonic pulses during Havel/Elbe deposition were associated with

Figure 4.3.2. Palaeogeography and general facies distribution of the Upper Rotliegend II Elbe Subgroup (compiled from Ziegler 1982 and Glennie et al. 2003). NPB = Northern Permian Basin, SPB = Southern Permian Basin

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H. Stollhofen et al. Figure 4.3.3. Interference of marine pre-Zechstein ingressions with climatic variability

basaltic volcanism (Gebhardt et al. 1991). Altmark tectonic pulses are expressed by the development of transtensional and transpressive structures, including abundant soft sediment deformation features. Contemporaneous with the Altmark tectonic pulses, up to 3000 m (Hertle and Littke 2001; Henk 1992: 2000 m) of sediments were eroded in the SW German Saar-Nahe Basin during a phase of basin inversion. In this context we assume Upper Rotliegend II phases of enhanced heat flow, erosion and development of unconformities in areas fringing the Central European Basin as a result of thermal uplift of the crust prior to successive rifting events that preceded the opening of the Tethys and the Arctic-North Atlantic Oceans. Tectonic controls on deposition Altmark I to IV tectonic movements triggered the formation of successive fining-upward units (Hoffmann et al. 1989; Hoffmann 1990; Gebhardt et al. 1991) represented by the Parchim, Mirow, Dethlingen and Hannover For-

mations (Plein 1995). Each of these formations covers about 2 to 3 Ma (Fig. 4.3.1), most probably reflecting third-order depositional cycles. In particular the aeolian sediments of the Havel Subgroup record effects of synsedimentary tectonics as they are solely preserved as syntectonic graben and half-graben fills (Drong et al. 1982). Following deposition of the Havel Subgroup, a period of tectonic quiescence favoured alluvial fan progradation and local peneplenation of fault-generated topography. This was succeeded by Altmark III tectonics which provided new accommodation space to deposit the Elbe Subgroup. Thicker salt accummulations are believed to be deposited only within (fault-bounded?) centres of subsidence such as the Glückstadt Graben, although the latter is regarded as a post-Permian structure (Maystrenko et al. 2005b). Climate and sea level control The two third-order Altmark III and IV tectonic cycles contain 14 nested fourth-order cycles (Fig. 4.3.1) of ~400 ka

Upper Rotliegend to Early Cretaceous basin development

duration (Gast 1995). Such cycles are reflected by variations in the level of the perennial saline lake that existed in the basin centre during Elbe Subgroup deposition (Gast 1991). Each highstand resulted in a considerable deepening and enlargement of the saline lake followed by lake regression. Claystones were deposited during lake level highstands, and during lowstands halite deposition prevailed in the basin depocentre, whereas lake margin areas developed extensive sand and mudflats on which small aeolian dunes migrated. Due to the low relief of the lake margins, even small lake level fluctuations resulted in considerable lateral shifts of facies. The lake level highstands are most probably linked to wet periods with enhanced precipitation that correspond to Milankovitch-driven global climate fluctuations of ~400 ka eccentricity cycles (Gast 1991; Yang and Baumfalk 1994). There is an additional depositional pattern occurring at a periodicity of ~1 Ma. These are the “P”, the “Ameland”and the “Bahnsen”-highstands representing the most pronounced lake level highstands and short-lived marine intervals such as the “Garlstorf”, “Niendorf” and “Munster” ingressions. They can probably be explained with climatic precession modulation, but effects of superimposition of different Milankovitch cycles cannot be excluded. It is remarkable that this periodicity can also be correlated with sea level highstands in the Arctic rift system (Fig. 4.3.1), and time-stratigraphic equivalent deposits in Eastern Greenland (Wegener-Halvø/Ravnefjeld Formations) show a similar cyclicity (Stemmerik 2001). Although Upper Rotliegend II deposition took place in a hydrologically closed, continental basin, short-termed marine ingressions occurred in advance of complete flooding of the CEBS by the major Zechstein transgression. However, lake horizons bearing marine fauna originating from such short-term marine ingressions cannot be distinguished from ordinary lake deposits solely on the basis of lithofacies investigations. The first ingression into the CEBS is known from the Garlstorf Member (Dethlingen Formation) and characterised by occurrences of Liebea reichei, a marine bivalve (Gast and Gebhardt 1995). A renewed marine influence occurred during deposition of the Niendorf and Munster Members (Hannover Formation). The Niendorf ingression is also reflected by the fossil record as the Zechstein fish Acentrophorus sp. together with Liebea reichei were found in Schleswig-Holstein (Reiche 1956). Liebea reichei is also known from wells in Brandenburg (Schneider and Gebhardt 1993). In addition, δ34S-values of anhydrite register the Niendorf and Munster ingressions (Legler 2006). In contrast to the Garlstorf ingression, which coincides with a period of increased precipitation, Niendorf and Munster ingressions are out of phase with climatic cyclicity (Figs. 4.3.1 and 4.3.3).

The pre-Zechstein marine ingressions used the same seaway for the flooding of the SPB as the subsequent Zechstein transgression did. This seaway extended from the Boreal via basins between Greenland and the Fenno-Scandian Shield, over the incipient Viking Graben into the Northern and Southern Permian Basins. Marine pre-Zechstein ingressions from the south can be excluded because Acentrophorus sp. and Liebea reichei both display a close affinity with the Arctic faunal assemblage. Pre-Zechstein ingressions resulted only in a shortterm influx of marine waters into the saline lake. Most probably, the CEBS basin floor had already subsided deeply below sea level (cf. van Wees et al. 2000) prior to Niendorf and Munster ingressions. Hence, a longer lasting connection with the marine realm, as caused by tectonic events, would have resulted in a complete flooding comparable to the Zechstein-transgression. Therefore, a model is favoured where marine ingressions were only possible during sea level highstands in the Arctic rift system (Fig. 4.3.3). Flooding of an assumed geomorphological barrier north of the Viking Graben was then achieved during relatively short-term events, such as heavy storm tides or even tsunamis. The Munster ingression, however, did not occur in phase with sea level highstands recorded in the Arctic rift system (Stemmerik 2001). Thus it may represent a potential example of an ingression forced by tectonic movements.

4.3.3 Zechstein Stratigraphic framework The Zechstein Group is characterised by various types of evaporites and starts with the Kupferschiefer, representing the first widespread marine sediment that followed the dominantly terrestrial Rotliegend Group. A Re-Os age of 257.3±1.6 Ma for the Kupferschiefer (Brauns et al. 2003) is consistent with the Wuchiapingian age of Kupferschiefer-equivalents in the Southern North Sea based on occurrences of the conodont Mesogondolella britannica (Legler et al. 2005). The Zechstein succession consists of stacked, more or less complete evaporation cycles (Fig. 4.3.4) referred to as the Werra (Z1), Staßfurt (Z2), Leine (Z3), Aller (Z4), Ohre (Z5) and Friesland (Z6) Formations (Subkommission Perm-Trias 1993; Kiersnowski et al. 1995; Käding 2000; Warren this volume). Two additional cycles, Z7 and Z8, are only rarely developed and are usually dominated by clays and sandstones.

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H. Stollhofen et al. Figure 4.3.4. Zechstein cycles, lithostratigraphy and sequence stratigraphy (modified from Kaiser et al. 2003). Left column shows German Zechstein cycles (Z1-Z8) (modified from Käding 1978; Subkommission Perm-Trias 1993; Kiersnowski et al. 1995). Columns to the right compare the sequence stratigraphic interpretations of the Zechstein in the Southern Permian Basin (SPB) (Strohmenger et al. 1996) and in northeast England and the adjoining North Sea (Tucker 1991)

Palaeogeography and depositional environments Zechstein strata reach thicknesses of about 1000 m and 1500 m in the Northern and Southern Permian Basins, respectively. Original depositional thicknesses of the Zechstein Group, however, are difficult to ascertain due to halotectonic movements, subsurface leaching and erosion at unconformities. Compared to the Upper Rotliegend II, the depositional area of the Zechstein reflects considerable growth in terms of its N-S extent. As the evaporation cycles developed in both the Northern and Southern Permian Basins, the hydrological separation between both sub-basins by the topographic barrier Mid North Sea High and Ringkøbing-Fyn High no longer existed (Taylor 1986). In addition, the CEBS newly incorporated the area of the Baltic Depression toward the NE and the Hessian Depression toward the S became considerably enlarged (Ziegler 1990; Becker and Bechstädt 2006). Apatite fission-track analyses in the area of the northern Rhenish Massif and the southern Ruhr Basin/Germany (Karg et al. 2005) as well as the recognition of a clastic Zechstein marginal facies in the Saar-Nahe Basin area (El Ouenjli

and Stapf 1995; Dittrich 1996) suggest that the southern limits of the Zechstein depositional area extended considerably farther south and southwest than shown on available palaeogeographic maps. At the beginning of the Zechstein, rifting in the ArcticNorth Atlantic region, probably contemporaneous with a global sea level rise, caused flooding of large parts of the CEBS through a seaway that connected the CEBS with the Arctic Seas (Smith and Taylor 1989). A basinwide ingression led to the deposition of the laminated, marly Kupferschiefer black shale in a euxinic sea where the bottom layer was stagnant and anoxic (Paul 2006a). Although the typical Kupferschiefer is only ~30 cm thick (Paul 2006b), this unit can be traced from England to Belorussia. Considering the regional distribution and the deposition of laminated deeper-water basal Zechstein Kupferschiefer below storm wave base, the Zechstein transgression can be regarded as a catastrophic event, causing a variety of soft-sediment deformation features in underlying dune deposits (Glennie 1989; Strömbäck and Howell 2002). Such rapid flooding (Smith 1979)

Upper Rotliegend to Early Cretaceous basin development

may be explained by a pronounced gradient toward the CEBS resulting from subsidence of the basin floor about 250-300 m below sea level (Glennie and Buller 1983; Fig. 4.1.3) or even more (van Wees et al. 2000) prior to ingression. The scenario compares well to the situation during the Messinian crisis, when the level of the Mediterranean Sea was drawn down by at least some 1500 m below the level of the world oceans (e.g., Clauzon et al. 1996). The Kupferschiefer is overlain by limestones which were deposited in a better aerated sea water column. During deposition of the overlying Zechstein successions, the former basin edges were progressively onlapped and finally overstepped as a result of the combined effects of global sea level rise and widespread thermal subsidence. Facies patterns were strongly influenced by the pre-existing relief over which the Zechstein Sea transgressed (Füchtbauer and Peryt 1980; Paul 2006b). Marine Zechstein faunas display a close affinity with the Arctic faunal province (Weyer 1979; Hollingworth and Pettigrew 1988). Zechstein microfaunas in Poland, however, also contain an admixture of Tethys-related species (Peryt and Peryt 1977) indicating at least temporary links with the Tethys. Arid climate in combination with eustatic sea level fluctuations and/or tectonic movements repeatedly restricted seawater influx and gave rise to the development of stacked evaporation cycles. Each cycle starts with marine clays, succeeded by carbonates, Ca-sulfates and rock salt, with the climax of evaporation being reached by precipitation of potash and magnesium salts (Peryt and Wagner 1998; Warren this volume). Carbonate-sulfate platforms, in places with steep, tectonically structured slopes (Kaiser et al. 2003) and microbial reefs, developed along the basin margins and on intrabasinal shoals passing basinward into thin organic carbonates and sulfates (Paul 1986, 2006b; Tucker 1991; Peryt et al. 1993). At least some of the carbonate platforms (e.g., Ca2, NE basin margin) are “rimmed platforms”, characterised by a narrow bar system and landward lagoons and coastal sabkha environments (Geluk 2000; Kaiser et al. 2003). The central and marginal parts of the platform provide sensitive records of sea level fluctuations as they develop extensive karstified and hiatal surfaces during sea level lowstands (Fig. 4.3.4). The Zechstein evaporitic system terminated with the deposition of “Bröckelschiefer” mud and sandstones (Z5 to Z8) in widespread, flat coastal sabkha environments. Geodynamics and basin tectonics Tectonic activity, referred to as the “Tubantian I event”, has been recorded in the basal Zechstein deposits of the Netherlands (Geluk 1999a; 2005) and more or less contemporaneous tectonic movements can be detected at the base of the Zechstein in the southern North Sea (Legler et al. 2005). Although there is only evidence of minor syndepositional

faulting in the CEBS, basin subsidence cannot be solely explained by the effects of cooling and compaction following an Early Permian thermal anomaly (Ziegler 1990). The pronounced Zechstein-transgression was triggered by rifting processes in the Arctic-North Atlantic region combined with a glacio-eustatic sea level rise (Ziegler 1990). Most probably the Zechstein Sea entered the CEBS via an incipient Viking Graben, a seaway which had already channelled previous, but short-lived, marine incursions during the Upper Rotliegend II (Legler 2006). There is no evidence for the existence of the central North Sea Graben during the Zechstein and it is doubtful whether the Horn Graben subsided actively during this time. Although plutonic activity in the Oslo Graben continued during the Late Permian, evidence is lacking for Zechstein tectonic activity at its southern extension, in the area of the Bamble Trough (Sundvoll et al. 1990; Sundvoll and Larsen 1994). Rapid subsidence of the Polish (Wagner et al. 1978) and the North Danish sub-basins suggests that tensional stresses, probably related to early phases of crustal extension in the Tethys realm, reactivated the Tornquist-Teisseyre fracture system as early as during the Zechstein (Ziegler 1990). Faunal affinities also mirror centres of crustal extension and subsidence as they indicate both a prevailing exchange with the Arctic faunal province and admixture of Tethysrelated species in Poland, the latter probably via an incipient, fault-controlled East Carpathian Gate (Ziegler 1990). Through the Hessian Depression, the Zechstein Sea ingressed southwards deeply into the Variscan fold belt (Paul 1985; Becker and Bechstädt 2006) thereby tracing a Late Carboniferous-Early Permian fault system that transected the Variscan deformation front (Gast 1988). The latest Permian then involves the inception of a complex, multidirectional rift system that transected the Variscan Fold Belt and the CEBS. Tectonic controls on deposition In general, the reconstruction of subsidence patterns during the Zechstein involves major uncertainties due to intense salt diapirism during the Mesozoic and Cenozoic, and also because of leaching and erosion effects in areas that became uplifted during the Mesozoic rifting phases. Particularly facies and thickness pattern of the first Zechstein cycle Z1 were controlled by syndepositional extensional tectonics along the southern margin of the CEBS. There, up to 300 m of halite accumulated in local sub-basins that were bounded by reactivated Carboniferous-Permian fault systems (Ziegler 1990). Further examples of such salt basins have been identified in the Emsland-Krefeld embayment between the Brabant and Rhenish Massifs (Ziegler MA1989), in the northern part

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H. Stollhofen et al. Figure 4.3.5. Overview section of the Buntsandstein Group (modified from Lepper et al. 1998 and Röhling et al. 2002) with major unconformities and marine incursions. Not to scale. Radiogenic age of the Permian/Triassic boundary (252.6±0.2 Ma) according to Mundil et al. (2004), all other ages are astronomically calibrated (Bachmann and Kozur 2004)

of the Hessian Depression and also in northern and southwestern Poland (Wagner et al. 1981; Wagner 1991). Additional evidence of syndepositional deformation comes from reported occurrences of possibly tectonically triggered slumps, slides, >100 m long crumbling folds and debris flow breccia within the first Zechstein cycle Z1 (Peryt 1992; Paul 2006b). Climate and sea level control During the Late Permian, the CEBS was at latitudes comparable to the modern day North African-Arabian deserts, 10–15° north of the equator (Glennie 1983). Aridity during deposition of the Zechstein Group is geologically evidenced by the development of stacked evaporation cycles and a drastic reduction of clastic input resulting from low precipitation. Modelling of Late Permian climates reveals a pronounced impact of monsoonal circulation pattern on the Pangea supercontinent (Kutzbach and Gallimore 1989; Parrish 1993). Well developed Zechstein evaporite cycles provide a convincing record of global sea level fluctuations which controlled sea water influx into the CEBS, most probably channelled by narrow, fault-controlled seaways in the areas of the incipient Viking Graben toward the N and the nascent East Carpathian Gate to the SE. These sea level fluctuations and resulting unconformities that formed during lowstands provide the basis for the Zechstein sequence stratigraphic framework (Tucker 1991; Strohmenger et al. 1996). It remains questionable whether the sea level fluctuations have

a glacio-eustatic origin as the main phase of Carboniferous-Permian Gondwana glaciation had already terminated at ~288 Ma (Bangert et al. 1999; Stollhofen et al. 2008). Although the third-order cycles (Z1-Z6) are, to some extent, very unequal in terms of thickness and number of lithological subunits their accumulation most probably took a similar amount of time (~1 Ma) if a total duration of about 7 Ma for the Zechstein Group is considered (Menning 1995). For nested small-scale cycles Richter-Bernburg (1985) inferred a “Milankovitch band” cycle duration of about 50-100 ka, based on counting the different varve laminae in the Zechstein carbonate-anhydrite laminites. Most time, however, might be represented by the unconformities separating the cycles (Richter-Bernburg 1950).

4.3.4 Buntsandstein Stratigraphic framework In the German part of the CEBS the Buntsandstein represents a predominantly terrestrial redbed sequence sandwiched between the marine Zechstein and Muschelkalk. The lithostratigraphic subdivision of the Buntsandstein Group into seven formations reflects stacked fining-upward cycles (Lepper and Röhling 1998), some of which are overlying basal unconformities (Fig. 4.3.5). The Zechstein/Buntsandstein boundary is considered as conformable and characterised by high influx of sand (Best 1989) with the first thick sandstone units defining the base

Upper Rotliegend to Early Cretaceous basin development

of the Buntsandstein. Biostratigraphy of the Buntsandstein is constrained by conchostracans (e.g., Bachmann and Kozur 2004). Although there are no radiogenic ages available to calibrate the Zechstein/Buntsandstein boundary directly, the U/Pb zircon age of 252.6±0.2 Ma (Mundil et al. 2004) for the Permian/Triassic boundary can be taken as a minimum value. This biostratigraphic boundary, as defined at the Meishan GSSP (Global Stratotype Section and Point) in China, is some 20 m above the lithostratigraphic Zechstein/Buntsandstein boundary. The boundary interval is constrained by the Tethyan marine scale using biostratigraphy, magnetostratigraphy and isotope stratigraphy (e.g., Bachmann and Kozur 2004), and its magnetostratigraphic correlation with sections of the CEBS in Germany is well established (Szurlies et al. 2003; Szurlies 2007). Palaeogeography and depositional environments The Buntsandstein extends W-E over more than 1500 km from Great Britain to Poland and N-S from Scandinavia to Switzerland. There is a pronounced NW-SE facies transition as fluvio-aeolian palaeoenvironments dominated e.g., in the East Irish Sea and Cheshire Basins, UK (e.g., Cowan 1993; Mountney and Thompson 2002) whereas the Alpine Buntsandstein of the Drauzug in Carinthia and eastern Tyrol, Austria was deposited at the margin of the Tethys Sea and comprises a dominantly fluvial to shallow marine facies (Niedermayr 1985; Stingl 1989; Krainer 1993). The depositional area of the CEBS is located between these two extremes and involves strata which record a spectrum of both, sporadic marine and aeolian influenced terrestrial depositional settings. The German Buntsandstein consists mainly of continental clastic deposits of sand seas, rivers and lakes with some calcareous and evaporitic units (Grumbt 1974; Clemmensen 1979; Paul 1982; Wycisk 1984; Olsen 1988; Radies et al. 2005). With the beginning of the Buntsandstein, the influx of clastic sediment into the CEBS increased, most probably due to a tectonic-induced relief rejuvenation in combination with higher precipitation. Sediment input from the north was efficiently trapped by the Danish and Polish sub-basins and only minor parts reached the North German sub-basin. In contrast, huge ephemeral river systems transported coarse clastics from the Massif Central and the Vindelician High northward via the Trier Embayment and the Hessian Depression, toward the Southern Permian Basin margins (Lepper and Röhling 1998; Paul et al. in press). There, braided fluvial sandstone successions dominate the Buntsandstein strata and interfinger basinward with cyclically arranged predominantly red, lacustrine mudstones and thin oolitic carbonates deposited in playa systems. Widespread “violet horizons” (VH), interpreted as palaeosols, are particularly well developed in southern basin margin areas and are used as stratigraphic markers (Ortlam 1974).

Geodynamics and basin tectonics Throughout the deposition of the Buntsandstein, the CEBS subdivided into several mostly NW-SE trending depocentres (e.g., southern North Sea, Danish, Polish, North German and South German sub-basins), some of which show an obvious association with faults (e.g., North Danish sub-basin: Tornquist Zone) or bordering “swells” (e.g., North German sub-Basin: Brandenburg High; Lepper and Röhling 1998). Another set of graben systems and “depressions” (e.g., Viking-Central-, Glückstadt, Horn and Westdorf Grabens; Hessian and Burgundy Depressions) as well as “swells” (e.g., Eichsfeld-Altmark, Hunte and Hannover Swells) are arranged obliquely to the long-axes of the sub-basins outlined. These structures reflect the inception of a complex, multidirectional rift system which is intimately linked to the interference of the rapid southward propagation of the Norwegian-Greenland Sea Rift and the westward propagating Tethys rift system (Ziegler 1990). In this context “swells” and “highs” represent the uplifted rift shoulders of adjacent grabens (cf. Frisch and Kockel 1997; Geluk 1999b; Kockel 2002). Buntsandstein thicknesses thus vary considerably between about 500 and 1200 m, but reach maximum values of up to 4000 m in local grabens in NW Germany and the southern North Sea (Lepper and Röhling 1998; Geluk 2005). Seismic, borehole and outcrop data across “swells” and “depressions” reveal contrasting thickness and facies developments, suggesting that these palaeotectonic and palaeogeographic entities were actively forming during the Buntsandstein (Bachmann and Grosse 1989; Röhling 1991; Lepper and Röhling 1998; Radies et al. 2005), at least since the Middle Buntsandstein (Hardegsen) tectonic phase (Frisch and Kockel 1997; Brückner-Röhling and Röhling 1998). This phase comprises several rifting pulses, four of which are of regional extent and recorded by the Volpriehausen (V), Detfurth (D), Hardegsen (H), and Röt (R) unconformities (Fig. 4.3.5) (cf. Geluk and Röhling 1999; Geluk 1999b). The base Solling (Hardegsen) unconformity can be traced from the Irish Sea and North Sea Basins as far as southern Germany (Fig. 4.3.6), illustrating a regional peak in differential subsidence (cf. Geluk and Röhling 1997). Subsidence patterns during the Buntsandstein are controlled by relatively short-lived WNW-ESE extensional rifting pulses with long-lasting phases of thermal subsidence in between (Brückner-Röhling and Röhling 1998). Tectonic controls on deposition The deposition of the Buntsandstein took place in NNW-SSE to NE-SW trending grabens and halfgrabens indicated by asymmetric infills, particularly obvious on the subcrop map of the base Solling (Hardegsen) unconformity (Fig. 4.3.6; Geluk and Röhling 1997). Direct evidence of synsedimentary fault control

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H. Stollhofen et al. Figure 4.3.6. Subcrop map of the base Solling (Hardegsen) unconformity (modified after Geluk and Röhling 1997)

in the Middle Buntsandstein is given by abrupt thickness changes in 3D seismic data sets and the observation of clastic dykes, convolute lamination and intraformational breccia in drillcores e.g., from the western margin of the Ems Trough (Radies et al. 2005). Outcrops of the Thuringian sub-basin illustrate a whole spectrum of subseismic-scale synsedimentary tectonic deformation structures, including normal and reverse faults and dm-scale mud diapirs in the basal Volpriehausen Formation (Heinzelmann 1969; Schüler et al. 1989; Voigt and Gaupp 2000). Further tectonic influence on sedimentation is indicated by strong thickness contrasts of the Middle Buntsandstein over short distances at the western margin of the Eichsfeld-Altmark High (Wycisk 1984) and the Glückstadt Graben (Fig. 4.3.6).

Indicated by the structural orientation (see above), this tectonic control on sedimentation was most probably caused by the reactivation of Late Palaeozoic lineaments (cf. Arthaud and Matte 1977). Regional tectonic movements related to extensional faulting in the CEBS triggered Zechstein salt movement during Middle Buntsandstein, ultimately leading to rafting of Buntsandstein blocks in the central parts of the subbasins and within grabens (Mohr et al. 2005). Cumulative effects of regional extensional faulting and salt movements were recorded by differential sedimentary facies and thickness evolution during the deposition of the Hardegsen Formation (Strunck et al. 1998).

Upper Rotliegend to Early Cretaceous basin development

Climate and sea level In principal, arid climate and the regressive trend of the uppermost Zechstein persisted into the Buntsandstein. Stacked fining-upward cycles, 10-30 m thick, are interpreted to reflect alternating humid and arid periods, controlled by 100 ka Milankovitch eccentricity cycles (Paul 1998). Although of mainly continental origin, the Buntsandstein in the CEBS was not deposited in a completely landlocked basin. Short-lived connections to the marine realm existed (Fig. 4.3.5) via the East Carpathian and the Silesian-Moravian Gates (Beutler and Szulc 1999) and possibly through the central North-Sea and Greenland Rifts (Michelsen and Clausen 2002). Bioturbated beds characterised by trace fossils of the Scoyenia and Cruziana ichnofacies in the Lower and Middle Buntsandstein (Lepper and Uchmann 1994; Knaust et al. 1999) and the Avicula Beds of the Middle Buntsandstein (Wycisk 1984; Radies and Tietze 1998) are stratigraphic levels that are discussed as equivalents of marine flooding surfaces (Lepper et al. 1998; Radies et al. 2005). During the early Anisian (uppermost Solling and lower Röt Formations) a pronounced transgression from the Tethys Sea, via the East Carpathian and the SilesianMoravian Gates, flooded much of the southern CEBS. This ingression favoured the accumulation of carbonates and sulfates in the lower Röt Formation of the eastern (Polish) part of the basin, and of sulfates and halites in the North German Basin and the northern parts of the Hessian Depression. A subsequent regression during the upper Röt did not affect these eastern parts of the basin, where carbonate deposition

persisted into the Anisian. During the uppermost Röt Formation widespread carbonates of the Lower Gogolian Beds in Poland and the Myophoria beds in Germany reflect a major sea level rise, heralding the Muschelkalk transgression which ultimately flooded large parts of the CEBS.

4.3.5 Muschelkalk Stratigraphic framework In contrast to the dominantly terrestrial clastic Buntsandstein red beds below and the clastic Keuper succession above, the Muschelkalk Group is dominated by marine limestones and marlstones. The Anisian to early Ladinian Muschelkalk Group subdivides into three subgroups (Lower, Middle, Upper Muschelkalk) which in turn comprise seven formations (Fig. 4.3.7) that are well developed in the North and South German basin facies. In large parts of Germany the base of the Muschelkalk Group is defined by the “Grenzgelbkalk”. In Poland, however, timestratigraphic equivalents (Lower Gogolin Beds) of the German uppermost Röt Formation (Buntsandstein) are included in the Lower Muschelkalk. No radiogenic dates are available for the Muschelkalk in the CEBS, but only for biostratigraphically calibrated successions of the northern Tethys shelf. There, Mundil et al. (1996) were able to obtain tight concordant clusters of single-zircon U-Pb results from tuff layers to estimate an age of 240.7-241.3 Ma for the Anisian/Ladinian boundary in Muschelkalk timestratigraphic equivalents of the Southern Alps.

Figure 4.3.7. Lithostratigraphy of the Muschelkalk Group in Germany (modified from Hagdorn et al. 1998). Age of the Anisian/Ladinian boundary according to Mundil et al. (1996), all other ages are astronomically calibrated (Bachmann and Kozur 2004)

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Palaeogeography and depositional environments The Muschelkalk Group attains a total thickness of 450 m in Poland and 300-500 m in northern Germany (Ziegler 1990). Middle Muschelkalk salts and thickness maxima mark the depocentre of the basin along the axis BerlinHamburg (Fig. 4.3.8). In principal Muschelkalk equivalent deposits show approximately the same extent as those of the Buntsandstein, but the depositional area probably advanced farther northwards to the northern rim of the Ringkøbing-Fyn High, due to the combined effects of sea level rise and enhanced subsidence of the North Sea Central Graben (Ziegler 1990). Palaeoenvironments reflect a shallow epicontinental sea that was mostly but not permanently connected with the Tethyan marine realm via the East Carpathian and Upper Silesian as well as the Burgundy Gates. Characteristic is a predominantly endemic marine fauna, particularly in the Upper Muschelkalk deposits. Whereas the eastern parts of the CEBS record prevailing open-marine conditions with low clastic input, Muschelkalk equivalents in Denmark grade northwards into alluvial-fan deposits (Ziegler 1990). The Lower Muschelkalk Subgroup is dominated by bioturbated, grey limestones and marlstones with a wavy texture and some tempestites, so-called “Wellenkalk” (i.e., wavy limestones). Seismites with typical lenticular and nodular fabric related to soft sediment deformation processes are relatively frequent (Szulc 1993, 2000; Knaust 2000). The monotonous nodular limestone facies is interrupted by three thin fossil-rich or oolitic limestone horizons termed “Oolithbänke”, “Terebratelbänke” and “Schaumkalkbänke” which form wide-spread marker horizons (Hagdorn et al. 1998). Maximum flooding is represented by an interval enclosing the ”Terebratelbänke“ which are characterised by abundant in-situ occurrences of the brachiopod Coenothyris “Terebratula” vulgaris in large parts of the CEBS (Aigner and Bachmann 1992). The regressive character of the uppermost part of the Lower Muschelkalk records deposition in lagoonal to coastal sabkha environments. A subsequent drop in sea level restricted communication with the Tethys Sea during the Middle Muschelkalk as evidenced by several evaporitic cycles (Brückner-Röhling 2000) consisting of dolomites, dolomitic mudstones, anhydrites and halites (Simon 1999; Brückner-Röhling and Heunisch 2004). The two lower evaporite cycles are topped by syngenetic dissolution surfaces, recording the influence of meteoric water as can also be seen from reduced δ13C values (Beutler and Szulc 1999). The Upper Muschelkalk is initially transgressive but then generally regressive and records the development of a large carbonate ramp system in the CEBS (Aigner

1985). Sections are characterised by cyclic alternations of fossil-rich limestone beds with intercalating tempestitic marlstones (Aigner 1985, 1999). Rich, predominantly endemic ceratite faunas allow a detailed biostratigraphic subdivision (Bachmann et al. 1999; Urlichs 1999). The most extensive transgressive event throughout the entire German Triassic is a shale-rich interval containing the so-called “Cycloides-Bank”, indicating basin-wide sediment starvation and the formation of firmgrounds (Aigner and Bachmann 1992; Fig. 4.1.3). In the top part of the Muschelkalk Group influx of clastic detritus increased in the eastern parts of the basin, resulting from uplift of the eastern and northern parts of the basin and closure of the Silesian-Carpathian Gates (Szulc 2000). Geodynamics and basin tectonics Evolution of the southern part of the CEBS was directly influenced by Tethys rifting through reactivated Variscan master faults which transmitted crustal motions from the Tethyan rift system to its northern periphery (Szulc 2000). With ongoing tectonic activity, the Tethys Basin itself experienced a progressive compartmentalisation into subbasins. Faunal migration from the Tethyan marine realm into its northern periphery generally followed the rift-controlled opening of seaways (gates) and thus provides a reliable record of tectonic activity. Open marine conditions developed particularly adjacent to the seaways along the southeastern basin margins, whereas westward and northward the environments became increasingly restricted. It was only during Upper Muschelkalk times that the circulation pattern reversed when the westward shift of the Tethyan spreading centre gave rise to the opening of the western (Burgundian) Gate and the eastern and northern parts of the basin became uplifted. As the North German and the North Sea basins were controlled by the North Atlantic-Arctic rift system, the central part of the CEBS was located in between these spreading centres and dominated by thermal subsidence (Szulc 2000). Tectonic controls on deposition Thickness and facies variations reflect continued differential subsidence of the major structural elements that transected the CEBS since the Buntsandstein. In particular the Horn, Glückstadt and Westdorf Grabens record enhanced depositional thicknesses during the Muschelkalk (BrücknerRöhling and Röhling 1998; Baldschuhn et al. 2001). Middle Muschelkalk halites fill a relief in the underlying strata (Aigner and Bachmann 1992) which is most probably faulting induced as indicated by extremely variable depositional thicknesses and the coincidence of isopach contours with structural elements (Fig. 4.3.8). An important effect of Middle Muschelkalk tectonic activity is the triggering of further rafting of Zechstein salt (Mohr et al. 2005; Fig. 4.1.3).

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Figure 4.3.8. Anisian-Ladinian Muschelkalk palaeogeography (modified from Ziegler 1990). MNH=Mid North Sea High, RFH=RingkøbingFyn High, ET=Emsland Trough, ND=North Danish Basin, PT=Polish Trough, WD=Weser Depression, TB=Trier Embayment, TW=ThuringianWest Brandenburg Depression, NL=North Netherlands Low, NN=North Netherlands Swell, RS=Rügen Swell, BS=North-East Brandenburg Swell, EA=Eichsfeld-Altmark Swell

New models of the genesis and environment of Lower Muschelkalk “Wellenkalk” based on studies in the Thuringian sub-basin relate a considerable part of the nodular textures to tectonic induced soft sediment deformation and place ruditic intraclastic limestones in a succession of earthquake induced seismites, debrites and turbidites (Knaust 2000). Some of these high-energy event layers can be correlated over 100-200 km. They are frequently associated with erosional structures (e.g., pot casts, gutter casts and channels), upper flow regime bedding features (e.g., antidunes) and soft sediment deformation structures (load casts, ball and pillow structures, convolute bedding, slide surfaces and slumps) grading laterally into debris or mud flows (Knaust 2000; Föhlisch and Voigt 2001). Further examples of seismically induced soft-sediment deformation have been described

from the Lower Muschelkalk of Poland and Germany (Szulc 1993; Rüffer 1996; Neuweiler et al. 1999; Knaust 2002). Climate and sea level control During deposition of the Muschelkalk Group, Central Europe was located at about 30 °N and experienced a subtropical climate (Parrish 1999). During the early Anisian, a relative sea level rise induced the Muschelkalk transgression and, as evidenced by faunal assemblages, re-established the connection between the Tethys and the southern CEBS via the East Carpathian and SilesianMoravian gateways (Fig. 4.3.8). Prevailing open-marine conditions in the eastern and central parts of the basin are reflected by the accumulation of Lower Muschelkalk car-

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bonates, containing a rich Tethyan fauna including sponge and coral reefs in Upper Silesia (Hagdorn et al. 1999). Both westwards and northwards these carbonates grade into dolomitic and evaporitic marls, with the whole system recording a westward facies shift and immigration of Tethyan fauna through time. At the beginning of the Middle Muschelkalk the East Carpathian Gate was closed and only restricted communication with the Tethyan realm existed via the Silesian Moravian Gate (Urlichs 1999). Middle Muschelkalk carbonates in Poland contain impoverished faunas, whereas evaporitic conditions prevailed in the western and central parts of the CEBS. As evidenced by the distribution of Middle Muschelkalk salts (Beutler and Szulc 1999) and Upper Muschelkalk faunal immigration patterns (Urlichs and Mundlos 1985), a new western seaway of the Tethyan realm was established through the Burgundian Gate and the extended Hessian Depressian (Hagdorn et al. 1998) during the Middle and Upper Muschelkalk. Improved communication with the Tethys Sea was then re-established at the onset of the late Ladinian via the southern parts of the Polish Trough and the then-existing Burgundy Trough and induced open-marine clear-water conditions. Muschelkalk strata show a pronounced cyclicity that has been interpreted in terms of orbital (Milankovitch-type) precession (~20 ka) and eccentricity (~100 ka) controlled sea level fluctuations (Götz 2004; Brückner-Röhling and Heunisch 2004; Bachmann and Kozur 2004). In particular the Latemar carbonate platform (W Dolomites, N Italy), located on the Middle Triassic shelf of the northwestern Tethys, has been presented as a prominent example of cyclic stratal patterns linked to orbital (Milankovitchtype) parameters (Hinnov and Goldhammer 1991). Some of these classic Milankovitch cycles, however, have recently been re-interpreted as much shorter, subMilankovitch fluctuations (Zühlke et al 2003; Emmerich et al. 2005).

4.3.6 Keuper Stratigraphic framework The marine conditions which prevailed during deposition of the Muschelkalk Group were replaced rapidly by the mostly non-marine environments of the late Ladinian to Rhaetian Keuper Group. This succession is subdivided into three subgroups comprising six formations (Fig. 4.3.9). The facies boundary to the marine carbonates of the Upper Muschelkalk is highly diachronous at the northern basin margin, whereas in southern Germany it is relatively sharp and characterised by a minor unconformity (Base Erfurt Unconformity D1), overlain by the

“Grenzbonebed”. Original radiogenic data are lacking in the Keuper strata of the CEBS; they are therefore imported from areas outside the CEBS and astronomically calibrated. Both Mundil et al. (1996) and Pálfy et al. (2003) dated Middle Triassic tuff beds in the Alpine and Hungarian Tethys at ~238 Ma that lie one ammonoid zone (~1 Ma) below the Carnian base. The base Carnian (~237 Ma) corresponds approximately to the base of the upper Grabfeld Formation (Estheria beds) (Fig. 4.3.9). Using the ~100 ka Milankovitch cyclicity of the lower Grabfeld, Erfurt and Meissner Formations, the base of the Keuper Group can be estimated at ~239 Ma. Palaeogeography and depositional environments In most parts of the Central European Basin (CEBS) the Keuper Group is about 400–600 m thick but shows considerable regional thickness variations, caused by the combined effects of rifting, transtensional tectonics and halotectonics. Compared to the Muschelkalk, the depositional area is slightly enlarged, particularly toward the NE and SW. Compared with Buntsandstein palaeogeography, the Lower Keuper reflects a complete reversal of the principal drainage pattern of the CEBS. Following closure of the East Carpathian and Silesian-Moravian Gates in southeast Poland (Szulc 2000), an extensive clastic, fluvial to deltaic/estuarine system prograded gradually southwards from Fennoscandia and interfingered with marine, restricted-marine and partly evaporitic, brackish marine and paralic environments along the southwestern basin margin (Wurster 1968; Paul et al. in press). Marine ingressions reached the basin exclusively through the Burgundy Gate and the Hessian Depression as reflected by faunal immigration patterns (Köppen 1997). The bivalve and ceratite bearing “Grenzdolomit” at the boundary of the Erfurt and Grabfeld Formations (Fig. 4.3.9) demarcates maximum flooding spreading as far north as central Germany (Aigner and Bachmann 1992; Köppen 1997; Pöppelreiter 1999; Szulc 1999, 2000; DSK 2005). During deposition of the Middle Keuper Grabfeld Formation the climate became generaly more arid and fluvial discharge almost totally ceased, resulting in widespread evaporites (Fig. 4.3.10). Sections consist mainly of clay- and siltstones with intercalated anhydrite beds and nodules, deposited in inland sabkha and playa environments (Cameron et al. 1992; DSK 2005). Several short-lived marine incursions entered the south-western CEBS through the Burgundy Gate. No direct marine connection has however been proved for the halite salinas in the North Sea or adjacent areas; the extent of a marine contribution to these salt accumulations (Halite A–E) is, therefore, uncertain. The basin wide pre-Schilfsandstein unconformity D2 (Figs. 4.3.9 to 4.3.11) resulted from

Upper Rotliegend to Early Cretaceous basin development

Figure 4.3.9. Stratigraphy, transgressive-regressive Depositional Cycles 5-6 (2nd order, Bachmann et al. this volume), 3rd order Depositional Cycles (I–IV) and unconformities (1–8) of the Keuper Group. Numerical, partly astronomically calibrated ages after Bachmann and Kozur (2004) and Kozur and Weems (2007). Location of Colnrade Z1 well is shown in Fig. 4.3.11. ECU=“Early Cimmerian” unconformity (base

a pronounced sea level lowstand prior to deposition of the Stuttgart Formation. This caused the formation of incised valleys (Aigner and Bachmann 1992), some as deep as 70-100 m in western and south-western Poland (Franz et al. 2005). At the eastern basin margin erosion may cut even deeper into the Lower Keuper and, locally, into the Muschelkalk. The Stuttgart Formation comprises fluvio-deltaic sandstones and mudstones, which spread over the entire CEBS (Wurster 1964; Beutler and Häusser 1982; DSK 2005). Palaeoenvironments resemble those of the Lower Keuper, but were dominated by fluvial activity. Only in the lowermost Stuttgart Formation is a restricted marine influence evident in Central and East Germany (Kannegieser and

Kozur 1972). Tidal influence has been recorded in Luxembourg and southern Germany (Barth et al. 1984; Gehrmann and Aigner 2002; Shukla and Bachmann 2007). During deposition of the Weser Formation, arid conditions returned and widespread sabkha mudflats favoured local halite formation. Halites accumulated at three distinct stratigraphic levels (Halites F–H) in isolated salinas that were smaller than those during deposition of the Grabfeld Formation (Fig. 4.1.2B). There were only a few marine ingressions into the CEBS, with the “Dolomie Beaumont” and equivalents (Fig. 4.3.9) in Southwest and North Germany representing the most widespread maximum flooding interval (Aigner and Bachmann 1992). Freshwater input fluctuated significantly as reflected by a few thin

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Figure 4.3.10. Correlation of the Erfurt, Grabfeld and Stuttgart Formations between the Central Graben and the Mid-Polish Trough, flattened at Base Stuttgart Unconformity (2). Horizontal distances not to scale

dolomite beds bearing freshwater to marine-brackish faunas. The “Lehrberg Beds” between the Halites F and G are the most widespread bundle of these dolomites and can be traced for more than 350 km across the CEBS. In the top part of the Weser Formation a stack of nodular gypcretes up to 50 m thick (Heldburg Gypsum) is widespread in Central and North Germany, Poland (Franz et al. 2007a), the Netherlands and offshore UK (Barnasch et al. 2007). The Arnstadt Formation was deposited in a playa system with a sparse limnic to brackish-water fauna (Will 1969). Characteristic of the whole formation is a pronounced cyclic arrangement of siltstones, marly mudstones and dolcretes that has been interpreted in terms of Milankovitch control (Reinhardt and Ricken 2000). Cyclic freshwater input is indicated by fluvial sandstones that prograded from the Bohemian-Vindelician Massif and the Fennoscandian Shield and interfingered with lacustrine and playa mudstones (DSK 2005; Franz et al. 2007b; Shukla et al. 2006). In large parts of the CEBS, the Upper Keuper Exter Formation records an environmental transition from Late Triassic non-marine environments, over paralic systems to Early Jurassic marine conditions. Connections with the open sea were situated farther west than during earlier Triassic times as new gates opened adjacent to the British Islands. There, the Tethys Ocean was probably connected with the CEBS through a new seaway between the Pennine High and the London-Brabant Massif (Ziegler 1990; Warrington and Ivimey-Cook 1992). The marine connec-

tion via the Burgundy Gate (DSK 2005) may also have persisted, whereas the East Carpathian Gate was most probably closed (Beutler and Szulc 1999). Toward the end of Keuper Group deposition the basin was almost entirely landlocked. Contrasting environments developed in the eastern and western parts of the CEBS. The western half of the basin was dominated by open marine deposits, mainly dark, grey to green bioturbated mudstones, occasionally rich in organic matter. Plant remains are common and suggest freshwater input (Will 1969; Cameron et al. 1992; DSK 2005). From bottom to top, there is a change from brackish (e.g., Unionites posterus Deffner and Fraas) to marine (e.g., Rhaetavicula contorta Portland) and finally limnic biota (mainly ostracods), recording major variations in the amount of marine influence. In contrast, the Upper Keuper in NE Germany and Poland consists entirely of nonmarine fluvio-lacustrine and sabkha deposits (Franz et al. 2007b). A N-S trending belt of shallow marine, coastal and deltaic sandstones with minor mudstone intercalations, approximately 100 km-wide, characterises the area where western marine and eastern non-marine conditions interfingered. The position of the belt remained remarkably stable in an area extending from Denmark to Lower Saxony and Thuringia (Germany). Marine influence diminished eastwards across this belt; the most easterly indications of marine influence are two thin dinocyst-bearing beds in southern Sweden (Lindström and Erlström 2006), and occasional occurrences of marine bivalves and limulids in central North Germany (Hauschke and Wilde 1991; DSK 2005).

Upper Rotliegend to Early Cretaceous basin development

Geodynamics and basin tectonics During deposition of the Lower Keuper, strong clastic sediment influx from northeastern sources dominated the depositional pattern (Wurster 1968; Ziegler 1990; Szulc 1999, 2000; DSK 2005). Fission track ages of detrital grains and K/Ar ages of micas indicate sediment provenance from the Norwegian Caledonides and the Oslo Graben (Köppen 1997; Köppen and Carter 2000; Paul et al. in press). These areas were the catchments of major fluvial systems which entered the north-eastern part of the CEBS through the Horn Graben east of the Ringkøbing-Fyn High (Beutler and Schubert 1987; DSK 2005). Similarly large fluvio-deltaic systems of the Middle Keuper Stuttgart Formation (Schilfsandstein) prograded from the north across the entire CEBS toward the Tethys (Wurster 1968). Rift shoulder uplift along rift basins between Northwest Europe and Greenland (Ziegler 1988, 1990), in combination with a temporarily more humid climate caused the southward transport of large quantities of clastics across the Fennoscandian Shield toward the CEBS (Kozur 1972, 1975; Simms and Ruffell 1989; Kozur and Weems 2007). The Keuper Group as a whole records a period of enhanced differential subsidence of the CEBS, most probably corresponding to an increasing activity of the Arctic North Atlantic Rift system (Brandner 1984; Ziegler 1990).

Associated tectonic movements are recorded by thickness and facies changes which become particularly obvious adjacent to swells (e.g., Eichsfeld-Altmark Swell), troughs (e.g., Glückstadt Graben) and at the basin margins. Another data source for the timing and amount of tectonic movements can be obtained from the study of unconformities in the sedimentary succession. Six unconformities have been identified in the German Keuper succession (Fig. 4.3.9) out of which the D2 and the D4 unconformities at the base of the Stuttgart and Arnstadt Formations (Figs. 4.3.11 and 4.3.12), respectively, are the most important ones (Beutler 1979). Establishment of the base Arnstadt (D4) unconformity, for instance, is associated with extensional tectonics in N-S striking zones (e.g., Glückstadt Graben), strike-slip movements of NW-SE striking structures (e.g., TornquistTeisseyre lineament) and also the initiation of salt tectonics in North Germany (Beutler 1998). The base Arnstadt (D4) unconformity, also referred to as “Early Cimmerian” unconformity thus records the most important tectonic event during the Late Triassic. Tectonic controls on deposition Three major extensional structures, all running NNE-SSW, are recognised during Keuper deposition: (a) the Horn-Ems Graben, (b) the Glückstadt Graben and (c) the Gifhorn rift zone (Frisch and Kockel 1999; Kockel 2002). Of these, the Glückstadt Graben records the most pronounced synsedi-

Figure 4.3.11. Subcrop map of the base Stuttgart unconformity (Unconformity 2) with important structural elements and location of Colnrade Z1 well. Abbreviations: BVM = Bohemian-Vindelician Massif, CG = Central Graben, DB = Danish Basin, EAS = Eichsfeld-Altmark Swell, EEP = East European Platform, EG = Ems Graben, FS = Fennoscandian Shield, FSM = Foresudetic Monocline, GG = Glückstadt Graben, LBM = London-Brabant Massif, MPT = Mid-Polish Trough, NS = Netherlands Swell, PH = Pennine High, RFH = Ringkøbing-Fyn High, RS = Rügen Swell, WBT = West Brandenburg Trough

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Figure 4.3.12. Subcrop map of the base Arnstadt (Early Cimmerian) unconformity

mentary subsidence, accommodating sediments of up to 5800 m cumulative thickness (Maystrenko et al. 2005b), whereas outside the graben thicknesses of only 400-600 m are common. In particular, evaporite deposition concentrated in the fault-induced depressions; therefore thick evaporite occurrences can be taken as a proxy for outlining synsedimentary active rift zones. Also basement faults such as the Uelzen, Aller and the southern Lower Saxony Lineament, and probably also the Osning Lineament, became reactivated with the beginning of sedimentation of the Grabfeld Formation. Mobilisation of Zechstein salt was triggered by rifting and caused salt pillows, salt diapirs and salt rafting (Trusheim 1957; Kockel 1998; Mohr et al. 2005; Fig. 4.1.3). The salt diapirs frequently straddle basement faults and often mark rift boundary faults. Late Ladinian and Carnian rift pulses initiated the main phase of salt diapirism in the CEBS, and therefore most salt plugs in the CEBS are of Keuper age (Jaritz 1973). Deep secondary rim synclines adjacent to salt diapirs may contain thick Keuper sediments and thus often conceal the true amount of tectonic rift subsidence, e.g., in the Glückstadt Graben. Locally, extension was so intense that the Lower to Middle Triassic cover of the Zechstein salts broke up into rafts and Keuper sediments were deposited directly on top of the Zechstein (Frisch and Kockel 2004). The Lower Keuper Erfurt Formation has a uniform thickness of ~80–125 m, suggesting a phase of minor tectonic activity in the CEBS, except the Glückstadt Graben,

where this unit attains up to 600-700 m thickness (Frisch and Kockel 1997, 1999). In contrast, strong extensional tectonics during the Middle Keuper Grabfeld Formation caused highly variable fault-controlled subsidence patterns. The usual maximum thickness of the Grabfeld Formation is <300 m. In the Glückstadt Graben, however, it is some 2700 m thick and in other NW German grabens it attains up to 1000-1500 m thickness (Geluk 2005). Enhanced thickness is mainly due to thick halite deposits that occur at up to five distinct levels and, despite their local fault controlled occurrence, can be correlated over large distances (Halite A-E; Frisch and Kockel 2004; DSK 2005; Geluk 2005; Barnasch et al. 2007). Fault-controlled differential subsidence during deposition of the Grabfeld Formation and subsequent erosion are expressed by the base Stuttgart unconformity D2, the relief of which is illustrated by a complicated subcrop pattern in figure 4.3.11. The Stuttgart Formation (Schilfsandstein) was then deposited during a period of relative tectonic quiescence and varies in thickness from 20 to >60 m. Average thicknesses of the overlying Weser Formation are 50-200 m but increase abruptly in synsedimentary graben structures such as the Ems Graben (<700 m) and the Glückstadt Graben (2500 m), where subsidence was partly caused by salt withdrawal resulting from salt diapirism. The base Arnstadt (Early Cimmerian) unconformity (Unconformity 4; Figs. 4.1.3, 4.3.9 and 4.3.12) caps the Weser Formation. This is one of the most pronounced Triassic

Upper Rotliegend to Early Cretaceous basin development

(Baldschuhn et al. 2001). Due to local uplift associated with the formation of Unconformity 6 the lowermost part of the Upper Keuper may be missing on swells (Netherlands Swell, Rügen Swell) and on basin margins (East European Platform). Climate and sea level control Climate during Keuper Group deposition was dominated by the northern reaches of the Pangaean megamonsoon at about 30-35° northern palaeolatitude (Golonka et al. 1994). It gradually changed from arid in the late Ladinian and Carnian, with a distinct middle Carnian humid interval (Kozur 1972, 1975; Simms and Ruffell 1989), to semiarid during late Norian and Rhaetian times. A pronounced cyclicity of sediment packages, some 10–30 m thick, is present throughout the succession but is most obvious in the sabkha mudstones and evaporites. Some of the cycles can be correlated for hundreds of kilometers and are thought to represent 100-400 ka Milankovitch-type climatic fluctuations, enabling the total duration of Keuper sedimentation to be estimated at ~37 Ma (Bachmann and Kozur 2004; Menning et al. 2005; Nitsch et al. 2005). Figure 4.3.13. Western Europe during the Early Jurassic, modified after Baudin et al. (1990)

unconformities in the entire CEBS and is especially evident at the uplifted northeastern, eastern and southwestern basin margins as well as on intra-basinal highs such as the Eichsfeld-Altmark Swell (Fig. 4.3.12) where erosion attains up to 500 m stratigraphic loss (Barnasch et al. 2005). However, only minor erosion is observed in most of the more central parts of the basin. The origin of the unconformity has been related to intraplate stresses exerted on the CEBS during the final closure of the Palaeo-Tethys (Ziegler 1990; Stampfli and Kozur 2006). Due to the relief of the base Arnstadt (Early Cimmerian) unconformity, the Arnstadt Formation and its equivalents rest on progressively older sediments toward uplifted areas (swells) (Fig. 4.3.12; Wolburg 1969; Beutler and Schüler 1978; Baldschuhn et al. 2001; Geluk 2005; Barnasch et al. 2005, 2007). The thickness of the Arnstadt Formation varies from ~50 m in northern Germany to ~200 m in the Central Graben reaching >400 m in the Glückstadt Graben. The formation terminates with the Base Rhaetian Unconformity 5. The Upper Keuper Exter Formation is on average 100-300 m thick in the North German Basin but thickens locally in graben structures, attaining values of up to 600 m in the southern part of the Glückstadt Graben

4.3.7 Jurassic Stratigraphic framework Beside the recent division of the Jurassic as detailed in the International Stratigraphic Chart (Gradstein et al. 2004), old, lithologically motivated epochs remained in use in the 20th century and sometimes are still being used, occasionally together with modern terminology within a single publication: Early Jurassic ~Lias ~“Schwarzer Jura”; Middle Jurassic ~Dogger ~“Brauner Jura”; Late Jurassic ~Malm ~“Weißer Jura”. At the stage level, quite a number of names were proposed, as discussed by Arkell (1933). Another complication arises because a shallow sea throughout Europe connected the Tethys (~28 °N) with the boreal realm in Greenland (~48 °N), allowing in part faunal migration. On the other hand, the shallow epicontinental sea was structured by islands and swells, causing bioprovincialism (Callomon 2003). For the latest stage of the Jurassic, therefore, there are still two terms in use, Tithonian and its boreal equivalent Volgian, both accepted due to remaining correlation problems. Recent comprehensive overviews about Jurassic bio- and chronostratigraphy are given by Page (2003, Early Jurassic), Callomon (2003, Middle Jurassic), and Zeiss (2003, Late Jurassic). For the Jurassic of Europe, a comprehensive overview

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H. Stollhofen et al. Figure 4.3.14. Jurassic thickness distribution in the CEBS (Courtesy Maystrenko, Bayer, Scheck). Structural elements: CG = Central Graben; GG = Glückstadt Graben; LSB = Lower Saxony Basin; MFB = Moray Firth Basin; NDB = Norwegian-Danish Basin; PB = Polish Basin; RVG = Roer Valley Graben; WNB = West Netherlands Basin; VG = Viking Graben

is given by Pienkowski and Schudack (coord., in press). For worldwide overviews, the books of Arkell (1956) and Hölder (1964) still provide reasonable sources, although stratigraphic terminology is not always up to date. The Triassic/Jurassic boundary has been assigned an age of 199.6±0.6 Ma in the International Stratigraphic Chart (Gradstein et al. 2004). Recent U/Pb dating (Schaltegger et al. 2008) and biostratigraphic investigations (Kozur and Weems 2007), however, suggest that the Triassic/ Jurassic boundary has an age of ~201.5 Ma. Basin differentiation, a major step towards complexity Following a first phase of basin reorganisation during the Triassic, with extensional tectonics, graben formation and associated salt tectonics (cf. Scheck-Wenderoth et al. this volume), Europe saw a major transgression and was covered by a large shallow epicontinental sea (Fig. 4.3.13). The onset of the transgression is already documented in the latest Triassic, the Rhaetian, with intervening intervals of terrestrial sediments and marine sandstones. The transgressive trend is particularly well documented by Triassic marine fossils and the famous “bone beds” containing teeth from marine as well as terrestrial vertebrates, and as an exception, of early mammals (e.g., Wagner 1960). At the beginning of the Jurassic, the sea prograded almost instantaneously northward and finally reached the CEBS. On the other hand, the Arctic seas may have prograded southward for the first time via the Rockall-Faeroe Trough Arctic faunas advanced southward through this seaway during the later Hettangian and reached the channel area via the Irish Sea where they interfingered with warm-water Tethyan fauna. The Jurassic transgression, however, was not spontaneous, resulting in an unresolved chronostratigraphic

resolution of the Triassic/Jurassic boundary in some parts of the CEBS. In the northernmost North Sea, the Netherlands on- and offshore, as well as in the Polish Trough the boundary is concordant but lying within coastal plain sediments that provide insufficient time control (Husmo et al. 2003; Dadlez 2003; Wong 2007). Therefore, lithological boundaries are frequently used, which may diverge from the chronostatigraphic boundary by several hundred metres. In other parts of the CEBS the transition is discontinuous; in parts of the basin system Middle Jurassic marine rocks overlie terrestrial Triassic unconformably. In addition to the pre-existing Triassic rift topography, the Jurassic succession within the CEBS was further complicated: Late Early to Middle Jurassic doming affected especially the central part of the North Sea, the Ringkøbing-Fyn High and most of the Danish Basin. This uplift caused widespread erosion in middle Aalenian to late Bajocian times and has been referred to as the“Mid-Cimmerian” unconformity. During the Late Jurassic the area was again under extensional stress, enforcing graben formation, especially in the North Sea, and block faulting. Stratigraphy becomes increasingly complex as Late Jurassic sediments are mostly preserved in the rift structures. The Late Jurassic extensional event reactivated salt movements wherever sufficient salt had remained. Salt rim synclines developed, hosting Jurassic sediments, whereas nearby the original sediment was eroded. In wide areas of the CEBS, the Jurassic-Cretaceous boundary coincides with a major unconformity, especially in the southern part of the basin, while in the northern part of the North Sea a continuous transition is observed. Consequently, Jurassic sediments have

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Upper Rotliegend to Early Cretaceous basin development

·

been removed locally and further complicate the sedimentary pattern. Not least, the late Cretaceous inversion (Voigt et al. this volume) modified the remaining Jurassic strata once more by additional erosion during uplift, especially along the basin margins and within the Polish Trough, hampering there the reconstruction of original thicknesses (Dadlez 2003).

In addition the palaeogeography of Europe changed, especially towards the end of the Jurassic. Regression became dominant as documented by the ‘fossil wood’ of Portland/ UK (former “Portlandian”), evaporitic to hypersaline conditions in the Lower Saxony Basin (LSB) or development of coral reefs in southern Germany. Finally, the seaway that had connected to the Tethys closed and the southern marginal seas of the CEBS vanished in wide areas. During the Jurassic euxinic conditions occurred repeatedly, some more or less locally, others on large to worldwide scales. The first thin black shale intercalations appeared in the Hettangian and Sinemurian, e.g., in the Netherlands and in southern Germany. During the early Toarcian, a major anoxic event during deposition of the “Posidonia shale” affected wide areas of Europe, providing a high class source rock. Within the basin area proper, the distribution is spotty due to later tectonics and erosion, as illustrated in Fig. 2.14 (Maystrenko et al. this volume). Another anoxic event is then observed in the Callovian (Peterborough Member, Oxford Clay Formation, e.g., Kenig et al. 2004). During the Late Jurassic, the Kimmeridge Clay Formation provided widespread first class source rocks in the central and northern North Sea, which together with the ongoing extensional tectonics and block faulting provided feeding of syn- and pre-syn-rift reservoirs (Fraser et al. 2003). A further anoxic event occurred at the Jurassic-Cretaceous boundary which can be considered a global Oceanic Anoxic Event (OAE, Langrock et al. 2002; Langrock and Stein 2004). Although it does not play a role within the CEBS, where by now shallow water conditions prevailed, it is widespread in the Norwegian Greenland Seaway, and it is recognised elsewhere at such distant regions as western Sibiria (Lopatin et al. 2003) and the Falkland Plateau in the South Atlantic (Ludwig et al. 1983). Due to the complex Jurassic evolution and the impact of subsequent events, the present-day distribution pattern of Jurassic sediments provides spotty patches in large areas of the CEBS together with strong thickness variations in areas where Jurassic sediments are more continuously preserved. Fig. 4.3.14 provides a rather large and crude pattern, scaling down will reveal increasingly spottiness. In summary, the complex tectonic history during Jurassic times together with a suite of actual and potential

source rocks prohibits any general hydrocarbon play for the entire basin system. Evaluation and production have to consider regional to local conditions in terms of sedimentology, tectonics and associated salt movement, providing a high risk for developing potential fields. Palaeogeography and depositional environments The sedimentary history of the CEBS throughout the Jurassic provides a rather complicated structure which can be elucidated here only in a rather general manner or with regard to patches of local developments. Concerning details, the reader is referred to the monographs of Evans et al. (2003) concerning the middle and northern North Sea, Ineson and Surlyk (2003) concerning the Danish part of the CEBS up to Greenland, Wong et al. (2007) for the Netherlands area, Boigk (1981), Betz et al. (1987) and Baldschuhn et al. (2001) for the German part, Dadlez et al. (1995) and Dadlez (2003) for the Polish Basin (Trough). Four major provinces have at least to be distinguished, (i) the northern North Sea, its (ii) central and (iii) southern portions, and (iv) the German and Polish Basin – the latter showing partly special conditions and an internal subdivision into a Kuiavian (central Poland) and Pomeranian (northern Poland) part (Dadlez 2003). (i) In the northern segment of the North Sea, the latest Triassic to Aalenian has been divided into three megasequences (Husmo et al. 2003): The Late Triassic to Sinemurian Statfjord Group, the Pliensbachian megasequence (consisting of a progradational upward-coarsening sandstone in the east intercalating with marine shales in the west which can be further subdivided into higher order cycles) and a Toarcian to Aalenian megasequence, with three upward-coarsening sandstone tongues representing maximum regression. The early Bajocian to early Bathonian is locally dominated by a progradational delta complex; the Bathonian and Callovian sequence then is dominated by sandstones, siltstones and shales. Extensional tectonics caused tilting of fault blocks in some parts and erosion of older strata. In addition, a volcanic episode occurred in the Callovian, feeding alkalic, olivine-basaltic magmas. The Late Jurassic sequence comprises diachronous lithostratigraphic formations like the Kimmeridge clay and sandstone dominated intervals, providing an interesting hydrocarbon play (Fraser et al. 2003). (ii) Ineson and Surlyk (2003) divide the stratigraphic development in Denmark (including the Central Graben) as well as in Greenland into three megasequences. The first, Late Triassic to Aalenian, is dominated by relative uniform subsidence and absence of major faulting. The

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H. Stollhofen et al. Figure 4.3.15. A) The NGB in the Early Jurassic (after Geyer and Gwinner 1986; Ziegler 1990). B) Palaeogeography and facies in the Hettangian-early Sinemurian. White areas are showing shallow marine environments. LB=LondonBrabant Massif, V=Rhenish Massif, B-V=Bohemian Vindelician High, FS=Fenno-Scandian High. C) Palaeogeography and facies in the Aalenian. Black arrows indicate transport directions, black dots in B and C locate occurrences of iron-oolites. In the NGB proper iron-oolites occur on top of structural highs. Gray arrows indicate Bajocian (S: sowerby zone) to Bathionan (B) sand transport. After Bayer (1989) and Brand and Hoffmann (1963)

sedimentary record however is fragmentary. EarlyJurassic rocks are missing in large parts of the area including the Central Graben, due to erosion during Late to Middle Jurassic uplift. Major uplift in the Central North Sea took place in late Early to early Middle Jurassic, separating the northern and southern parts of the North Sea for some time. The resulting unconformity within the North Sea is more or less constrained to the middle Aalenian.

During the subsequent syn-rift megasequence deposition was confined to narrow subsiding grabens in the late Aalenian to Bajocian. The rift climax then occurred during the Oxfordian to Tithonian (Volgian). Towards the end of the Jurassic, rifting diminished. Early Triassic to middle Aalenian sediments are mostly missing in the Danish part of the Central Graben. In the Danish Basin, as far as preserved, they consist of offshore basinal mudstones with intercalating shallow marine and

Upper Rotliegend to Early Cretaceous basin development

lagoonal sandstones. From middle Aalenian to late Callovian, flood plain mudstones with intercalating fluvial and estuarine sandstones dominate in the Central Graben. In contrast, basinal mudstones, fluvial and shallow marine sandstones dominate the Danish Basin, however, interrupted by several major unconformities. This sedimentary pattern continues into the Early Cretaceous within the Danish Basin and the Central Graben, where offshore mudstones dominated since the early Oxfordian. (iii) In the Netherlands section of the southern North Sea (southern Central Graben to onshore Netherlands), three depositional successions have been defined (Herngreen et al. 2003): Megacycle I extends from latest Triassic (Rhaetian) to Aalenian. It consists of open marine clays with intercalating limestone layers and bituminous shales of the Hettangian, Sinemurian, and last of all the Toarcian “Posidonia shale”, providing a first class source rock. Two unconformities are recognised

regionally, in the late Pliensbachian and late Toarcian, being also recorded in the southern part of the CEBS, e.g., in southern Germany. Megasequence II lasted from Aalenian to Callovian/ Oxfordian, providing important structural differentiation (“Mid-Cimmerian unconformity”) related to the development of the Central North Sea dome. Following Wong (2007), sedimentation again was dominated by open marine clays with intercalations of coastal sands in the late Aalenian. This sedimentation was interrupted in the Early to middle Bajocian by coastal sands, in places associated with local unconformities. Around the Bajocian/Bathonian boundary the sediment pattern changed again, with lacustrine restricted carbonates now interfingering with lagoonal clay- and siltstones. Megasequence III is floored by a diachronous disconformity starting during the early Callovian in the Broad

Figure 4.3.16. Kimmeridgian-Tithonian palaeogeography and facies distribution after Ziegler (1990). BM: Bohemian Massif, LBM: LondonBrabant Massif, LH: Lusatian High, PS: Pompeckj-Swell (Block), RFH: Ringkøbing-Fyn High, RM: Rhenish Massif

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Fourteens Basin, during the late Callovian in the west Netherland Basin, and during the Oxfordian in the Roer Valley Graben (Wong 2007). The overlying sedimentary sequence, continuing into the earliest Cretaceous, then consists mainly of fine grained coastal plain and lacustrine sediments intercalating with coastal sands and being interrupted by several unconformities. (iv) The North German Basin (NGB) and the Polish Trough were less or almost not affected by the North Sea Dome and the subsequent extension. An exception is the Lower Saxony Basin, which came into existence as a sub-basin during the extensional phase together with the Central Netherlands Basin which can be understood as its western continuation (Wong 2007). During the Hettangian to Aalenian the supply of clastic material from the Baltic area, the East European Platform and, perhaps, from the Bohemian-Vindelician High controlled sedimentation patterns. Within the NGB marine shales interfingered with lacustrine and sandy layers from the east (Boigk 1981). In parts of the North-East German Basin and especially in Poland, the Early Jurassic is dominated by non-marine deposits (Dadlez 2003). During the late Hettangian a sandstone facies belt developed, with its western margin following a north-south direction from Hamburg to Hannover and further south to southern Germany via the Hessian Depression, separating the Rhenish Massif from the Bohemian-Vindelician High (Fig. 4.3.15). Together with reworked horizons and iron oolite layers in the Sinemurian, the sequence represents a first coarsen-

ing upward sequence. The “Angulaten Sandstone” can be interpreted as coastal sandbars resulting from north-south directed currents. The Sinemurian to Pliensbachian consists of marine shales with some coarse grained intervals in the northwestern part of the basin, similar to the Central Netherlands Basin. The eastern part is more sandy and limnic. In the early Toarcian the “Posidonia black shales” were deposited (see Fig. 2.14), grading towards Poland into brackish to lacustrine sediments. During the Aalenian marine conditions approached the Kuiavian part of Poland while in the Pommeranian part marine conditions were established only during the Late Bajocian. In the NGB a situation developed similar to the Hettangian: sediment supply was mainly from the east, and coastal sands occurred along a north-south trending line from about Lübeck to the western margins of the Harz Mountains (Boigk 1981). The latter were spreading once more into the shallow sea of southern Germany, indicating a north-south directed current system. Several sandstones layers (sinon-, sinon-staufensis-, obtusa- and concave-sandstone) provide some cyclicity which can be correlated with sedimentary cycles and faunal overturns in the south German sea (Bayer and McGhee 1984, 1985). This situation may have continued into the middle Bajocian (Fig. 4.3.15c). During the late Middle Jurassic marine conditions in Poland prevailed and the general sedimentation pattern in Figure 4.3.17. Third order depositional cycles in the western Saxony basin compared with Great Britain and northern Switzerland. From Kästner et al. (in press)

Upper Rotliegend to Early Cretaceous basin development

the NGB changed. Sediment supply was coming from the North providing marine shales with intercalating sands that had their most southward extension in the Bathonian with a termination north of the Hessian Depression. The connection to the south was restricted, at least for sediment transport. In south-west Germany, Luxembourg and Lorraine a carbonate platform continuously built up which originated in the famous Minette oolitic iron ores (Luxembourg, Lorraine) during the latest Toarcian to Aalenian. Carbonate platforms in England show a similar development, indicating a major change in sedimentary conditions. At the beginning of the Late Jurassic, the palaeogeographic setting of the NGB drastically changed (Fig. 4.3.16; Boigk 1981; Betz et al. 1987). The connections to the southern shallow seas closed, and parts of the original Southern Permian Basin became uplifted, with the Pompeckj Swell limiting the east-west trending basin to the north. The NGB then was more or less restricted to the Lower Saxony Basin with connections to the Netherlands and Poland. Carbonate rich sediments dominate the Lower Saxony Basin and the Kuiavian part of Poland where a connection to the Tethys may have existed. In the Pomerian part, clastic sedimentation continued. The sedimentation within the Lower Saxony Basin resembled the situation in the Central Netherlands Basin, with some local specifics. During the Late Jurassic, the Lower Saxony Basin became subdivided into small troughs with lateral thickness and facies variations indicating the development of horst and graben structures (Betz et al. 1987; Gramann et al. 1997). Callovian and early Oxfordian consist of a shaly mudstone sequence of the Ornatenton Formation. The early to middle Oxfordian is characterised by bioturbated marlstones with variable amounts of sponge spicules and quartz grains. The mass appearance of sponge spicules is also known from South England (Mock and Palmer 1991; Wright 1992) and Switzerland. In South Germany, buildups of siliceous sponge are common (Brachert 1992). The middle to late Oxfordian is characterised by a carbonate-dominated ramp facies of the Korallenoolith Formation. This formation is restricted to the eastern part of the Lower Saxony Basin and consists mainly of oolitic limestones, iron-oolites, and micritic limestones with variable amounts of quartz sand. Coral-rich facies are restricted to a few horizons. Toward the west, the Korallenoolith Formation is missing, either due to replacement by shales and fine- to coarse-grained sandstones of the so called Wiehengebirgsquarzit (Klassen 1984) or because of nondeposition (Gramann et. al 1997). The vertical evolution of the Oxfordian deposits illustrates a gradual shallowing- and deepening-upward trend. The latter is indicated by a lithological change from bioturbated marlstones of the Heersum Formation deposited in a mid-ramp setting between storm wave base and fair-weather wave base into

shallow-water inner ramp swaly cross-stratified bioclastic and planar cross-stratified oolitic limestones of the lower Korallenoolith Formation. A temporary subaerial exposure with karst phenomena at the top of the lower Korallenoolith Formation records the regressive maximum. A major change in depositional style is observed above this major unconformity that is associated with an increase in quartz content (Kästner et al. in press). Deltaic siliciclastic deposits of the so-called Wiehengebirgsquarzit (Klassen 1984) were shed from the west into the basin, interfingering with carbonates of the upper Korallenoolith Formation. The vertical trend of the upper Korallenoolith Formation represents the deepening part indicated by the progressive change from inner-ramp quartz- and iron-rich oolitic limestones to deeper foraminifera-rich limestones of the mid-ramp setting. The larger-scale sequences are formed by a sequence set of five third-order sequences that can be correlated throughout the study area (Fig. 4.3.17). Correlation of third-order sequences show lateral thickness and facies variations of the Korallenoolith Formation indicating the development of a half-graben structure by synsedimentary tectonics (Gramann et al. 1997; Kästner et al. in press). This structure developed during the deposition of the lower Korallenoolith and was less intensified during the upper Korallenoolith Formation. Highest thickness is observed in the Wülpker Egge section with a progressive thinning towards the east and an abrupt decrease in thickness towards the west. Third-order cycles in the Oxfordian are widespread, occurring in England (Newell 2000), the eastern Paris Basin (Carpentier et al. 2007), and the Swiss Jura (Gygi et al. 1998); Newell (2000) and have been related to tectonic activity Carpentier et al. (2007). During the early Kimmeridgian, brackish-marine conditions dominated, soon giving way to fully marine conditions (Betz et al. 1987). Lateral thickness variations indicate a subdivision of the Lower Saxony Basin into horst and graben structures. The Kimmeridgian shales, sandstones and carbonates attain thicknesses of up to 200 m. The Tithonian finally is characterised by carbonate-anhydrite-halite cycles indicating increasingly restricted to evaporitic and hypersaline depositional conditions (Betz et al. 1987). Progressive shallowing resembles the evolution in the southern North Sea. Sedimentary thicknesses achieve 1000 to 1500 m. Rapid facies and thickness changes indicate tectonic movements along more or less North-South trending faults (Betz et al. 1987; Brand and Hoffmann 1963). Geodynamics and basin tectonics The Arctic-North Atlantic rift system remained active during the entire Jurassic and only during the Late Cretaceous evolved into the break-up axis of Laurasia (Ziegler 1990).

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During the Early Jurassic, rift-related magmatism was at a very low level but Mid-Jurassic uplift of the Central North Sea Rift dome in the Arctic-North Atlantic domain was accompanied by the development of a major volcanic complex and the emplacement of plateau basalts. The dimensions of this arch can be reconstructed by mapping the base late Bajocian (“Mid-Cimmerian”) unconformity over a distance of some 700 km in a north-south direction and ~1000 km east-west, the timing of its uplift being well recorded by the onset and waning of clastic supply from elevated areas (Brown 1984; Michelsen et al. 1987). Large parts of the North Sea area (e.g., Ringkøbing Fyn High) became affected by thermal uplift and erosion. Massive clastic input led to the interruption of open-marine seaways between the Arctic and Tethys domains, evidenced by faunal separation. In the NGB, clastics derived from the North Sea Dome form important reservoir rocks (Boigk 1981). Crustal separation between Africa and Laurasia during the Bajocian-Bathonian marked the onset of a new kinematic regime in the Atlantic-Tethys area (Ziegler 1990). The Callovian to early Oxfordian stress re-orientation is recorded by the abandonment of NE-SW trending grabens and troughs such as the Glückstadt and Horn Grabens, the Emsland and Gifhorn Troughs (Betz et al. 1987; Ziegler and van Hoorn 1989; Maystrenko et al. 2005b) and the development of new, NW-SE trending shear systems, eg. at the southern termination of the Central Graben of the North Sea rift (Ziegler and Dèzes 2006). Some of the latter reactivated Carboniferous-Permian fracture systems and controlled rapid subsidence of transtensional structures like the Sole Pit, Broad Fourteens, Central and West Netherlands and Lower Saxony Basins (Mazur and Scheck-Wenderoth 2005; Lohr et al. 2007a). These structures accommodated the ~15 km extension of the Central Graben during Late Jurassic to Early Cretaceous times. Contemporaneous to rifting, regional uplift (e.g., of the area around the Glückstadt Graben: Maystrenko et al. 2005b) and uplift of graben shoulders together with a global sea level lowstand caused subdivision of the CEBS into numerous sub-basins and small islands. Probably at the same time the long-lived seaways via the Trier Embayment and the Hessian Depression that had linked the CEB (North European Basin) with the Tethys shelf since the Mid-Triassic became closed (Ziegler 1990). Tectonic controls on deposition Isopachs of Early Jurassic strata show that basins continued to subside along patterns established during the Triassic. Quantitative subsidence analyses of the Southern Permian Basin indicate that its thermal subsidence continued during the Early and Middle Jurassic; the same probably also applies to the Northern Permian Basin (Sørensen 1985). Only the northwestern margin of the Southern Per-

mian Basin was affected by the Bajocian-early Bathonian thermal uplift associated with the emplacement of the Central North Sea Dome (Ziegler 1990). Following the Middle to Late Jurassic re-orientation of the stress regime and the development of new NW-SE trending shear systems, further evolution of the NW European rift systems was exclusively governed by stress regimes related to continued crustal extension across the Arctic-North Atlantic mega rift (Ziegler 1990). In the Lower Saxony Basin, the Jurassic fault pattern reflects activity of NW-trending faults, that developed as transtensional faults forming a pull-apart basin, due to strike-slip faulting at depth (Lohr et al. 2007a). In the centres of subsidence such as the Central Graben, Late Jurassic sediments reach a thickness of >1000 m, coupled with isostatic uplift of the adjacent footwall blocks on which the Late Jurassic strata are either missing or only thinly developed and interrupted by unconformities (Yielding 1990). Similar thickness is observed in the eastern Lower Saxony Basin and in the N-S striking Gifhorn Trough, but in the central Lower Saxony Basin the succession of Late Jurassic limestones, marls and evaporites is >2000 m thick. Inversion of this subbasin, especially in its western part, resulted in a nearly complete erosion of the basin-fill (see Voigt et al. this volume), and the reconstructed uplift of ~6000 m points to even higher primary thicknesses. Strong subsidence of the graben structures, uplift of horsts within the shear systems and block rotation at normal faults caused a complicated pattern of depocentres, unconformities (“Cimmerian unconformities”) and small-scale uplifts acting as local source areas. Accommodation space was rapidly filled with carbonates, so that nearly the whole Late Jurassic succession is dominated by oolitic limestones and lagoonal carbonates. Climate and sealevel control Quantitative sporomorph data from the southern Central Graben offshore The Netherlands indicate subtropical temperatures and humid conditions for the late Callovian and early Oxfordian (Abbink et al. 2001). Such climatic conditions at a time of high sea level favoured the growth of widespread carbonate ramps in Western and Central Europe (e.g., Newell 2000; Dupraz and Strasser 2002; Pawellek and Aigner 2003; Ruf et al. 2005). Thereafter, stepwise warming and aridisation occurred with maximum temperatures and aridity reached during the late Kimmeridgian and Portlandian (Abbink et al. 2001). Gypsum and halite deposits typically occur in latest Jurassic sediments (e.g., Münder Mergel) and reflect this trend. Multiple regressions and transgressions that are evident, particularly in the Middle Jurassic, indicate repeated

Upper Rotliegend to Early Cretaceous basin development Figure 4.3.18. Geographic map showing the areas referred to. I = Emsland, II = Teutoburger Wald, III = central basin, IV = Salzgitter area, V = Northern Subhercynian area, VI = Mecklenburg-Prignitz-Alt mark-Brandenburg area, VII = Usedom area (Mutterlose and Bornemann 2002)

sea level fluctuations which at first sight show no direct correlation with the global eustatic sealevel curve of Haq et al. (1987). The overall pattern of Jurassic sea level changes records a more or less gradual rise, interrupted by episodes of comparative stillstand rather than eustatic fall, and several episodes of significant regression have been identified to result from regional tectonics (Jacquin et al. 1998; Hallam 2001; Carpentier et al. 2007).

4.3.8 Early Cretaceous Stratigraphic framework The base of the Cretaceous is defined by the lowermost of three major Lower Berriasian to earliest Upper Berriasian unconformities (Late Cimmerian unconformities), in places developing an angular contact. It may truncate Jurassic strata, or even late Triassic sediments (Lohr et al. 2007a; Littke et al. this volume). According to the current International Stratigraphic Chart in use, the age of the Jurassic/ Cretaceous (=Tithonian/Berriasian) boundary has an age of 145.5±4.0 Ma (Gradstein et al. 2004). In the Lower Saxony Basin, the Jurassic/Cretaceous boundary has been placed within the “Münder Formation” and given an age of ~142 Ma (German Stratigraphic Commission 2002). The overlying “Bückeberg-Formation” is of Berriasian (Cretaceous) age. Marine conditions prevailing during the Jurassic were partly replaced by terrestrial de-positional environments and only in the Valanginian, did fully marine conditions return. The Early/Late Cretaceous boundary has been assigned an age of 99.6±0.9 Ma (Gradstein et al. 2004; German Stratigraphic Commission 2002).

Palaeogeography and depositional environments Earliest Cretaceous tectonic activity, combined with changes in sea level and climate, caused a termination of the late Jurassic carbonate-dominated deposition. Shallow marine carbonates were replaced by siliciclastic sediments common throughout the southernmost basins of the proto North Sea. These basins include the East Netherlands Basin, the Lower Saxony Basin, and the Danish-Polish Basin which had formed during the late Jurassic as faultbounded sub-basins of the CEBS south of the North Sea Central Graben and north of the London-Brabant-RhenishBohemian Massif. In the following, the stratigraphy and the sedimentary and tectonic record of one of these basins, the 80 x 280 km LSB, will be discussed as a type example. The LSB was bounded by the Pompeckj Swell to the north, by the Rhenish Massif to the south, by the East Netherlands High to the west and by the East Brandenburg High to the east (Fig. 4.3.18). In the western, southern and easternmost part of the LSB, shallow marine siliciclastics of Berriasian to Albian age reach thicknesses of up to several hundred metres (Fig. 4.3.19). This marginal facies interfingers with a basinal facies represented by dark mudstones up to 2000 m thick. The distribution and facies architecture of the sediments as well as their thickness are controlled by essentially three factors: differential subsidence, local tectonics and sea level changes. The latter are not only reflected by facies patterns but also by fossils of different biogeographic provenance with the following events being identified throughout the basin: (i) the Berriasian regressive phase, (ii) the early Valanginian transgression, (iii) the early late Valanginian transgression, (iv) the mid

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Figure 4.3.19. Lithostratigraphy of the Early Cretaceous successions (Berriasian–Albian) in the Emsland, the Teutoburger Wald, the Minden– Hannover–Braunschweig area, the Salzgitter area, the Northern Subhercynian area, the Mecklenburg–Prignitz–Altmark–Brandenburg area and the Usedom area. See Fig. 4.3.18 for locations. Modified from Mutterlose and Bornemann (2002)

Upper Rotliegend to Early Cretaceous basin development

Hauterivian transgression, (v) the Barremian regression, (vi) the early Aptian anoxic sediments, and (vii) the mid Albian hemipelagic marls. The Berriasian was a period of widespread regression, causing the isolation of several marginal marine basins in NW Europe. Throughout the LSB non-marine “Wealden” sediments were deposited under brackish-lacustrine conditions with several short-lived marine incursions well documented from the Bückeberg Formation. Sediments are up to 400 m thick along the western basin margin (Emsland area), but attain a thickness of 300-700 m further north and north-east in the basin centre (Minden– Hannover–Braunschweig area).

The early Aptian transgression caused significant changes in the palaeogeography. New sea-ways between the Tethyan and the Boreal Realms opened via northern France and southern England. In the LSB, Aptian strata achieve a thickness of at least 200 m and are dominated by clays and marls. Dark clays (bodei Clay) are overlain by the Fischschiefer (= Ocean Anoxic Event 1a), a finely laminated sediment rich in organic matter (Corg up to 7%), followed by stacked hemipelagic marls: the ewaldi Marl, the clava Marl and the inflexus Marl. The latest Aptian consists of dark clays (jacobi-nolani Clay). In the uppermost Aptian, sandy sediments up to 220 m thick (Rothenberg Sandstone) were deposited at the southwestern basin part, bordering the Rhenish Massif (Fig. 4.3.19).

A major transgression of the Tethys via the Carpathian sea-way and the return of marine conditions mark the basal Valanginian including the fully marine Platylenticeras Beds in NW Germany. The Gifthorn Channel, the Hoya Channel and the Ems Channel linked the LSB to the southern North Sea via the Pompeckj Block. At the western basin margin (Emsland area) dark grey claystones are up to 350 m thick, including several intercalated sandstone horizons (Bentheim Sandstone, Dichotomites Sandstone, Grenz Sandstone). The southern margin (Teutoburger Wald area) was flooded, except for the area situated farther southeast, where the onset of marine conditions occurred only during Hauterivian times. In the central part of the basin 300 m of marine clays were deposited during the Valanginian (Fig. 4.3.19).

The Albian saw yet another major expansion of the LSB depositional area. During the early Albian, black clays (schrammeni Clays) are extremely widespread, while pale marls were deposited during the mid and late Albian. The coast line was shifted quite considerably toward the south and southeast. During the late Albian the coastline was displaced even further to the south upon the Rhenish Massif. In the basin centre, Albian clays and marls, up to 250 m thick, are widespread. Further toward the east, the early Albian is represented by 40 m of glauconitic, sandy clays and sandstones (Hils Sandstone). The middle Albian consists of a 10-15-m-thick clay sequence (minimus Clay) and the late Albian consists of 80 m of biosiliceous marls (Flammenmergel; Fig. 4.3.19).

An early Hauterivian transgression further expanded the depositional area of the LSB with maximum flooding being recorded by the Endemoceras amblygonium ammonite Zone. A subsequent regression of late early Hauterivian age corresponds to the formation of the noricum Sandstone. The western basin margin (Emsland area) is dominated by up to 500 m thick medium grey clays with two sandstone horizons interleaved: the noricum Sandstone (mid early Hauterivian) and the Gildehaus Sandstone (late early Hauterivian) (Fig. 4.3.19).

As in the Late Jurassic, the Early Cretaceous evolution of the CEBS was dominated by crustal extension across the Arctic-North Atlantic rift system. Consequently, only few changes occurred in its structural framework and basin development (Ziegler 1990). These changes include major tectono-eustatic sea level variation and an overall increase in tectonic activity, corresponding to the Berriasian (“Late Cimmerian”) rifting pulse (Ziegler 1990). Persistent crustal extension of the North Sea rift resulted in a sharp accentuation of marginal troughs and uplift of the London-Brabant-Rhenish-Bohemian Massif. On a regional scale, rapid subsidence of several sub-basins (marginal troughs) occurred, leading to deep burial and high sedimentation rates (Bachmann et al. this volume, Littke et al. this volume), especially in the LSB. A likely source of the Aptian and Aptian/Albian boundary fallout tuffs, occurring in the area of Hannover, is probably the Waddenzee Volcanic Centre off The Netherlands (Mutterlose et al. 2003). In the Albian, however, tectonic activity in the North Sea Rift and the marginal troughs ended gradually and thermal subsidence began, leading to further cooling of the basin. The same applies to the Danish-Polish Trough (Ziegler 1989; Ziegler and Dèzes 2006).

A regression during the Early Barremian caused brackish-lacustrine conditions in central and southern Poland. In the LSB, however, finely laminated sediments, known as the Hauptblätterton, and various thin Blätterton horizons are typical of the Barremian. These sediments are enriched in organic matter (6-8% TOC) and were deposited under anoxic conditions, reflecting a stable stratification of the water column due to periodically increased run-off and related productivity changes. Along the western basin margin (Emsland area) Barremian sediments occur in the Brechte syncline where they attain a maximum thickness of up to 400 m.

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The Late Cretaceous and earliest Cenozoic evolution of Central Europe reflects a fundamental change in its megatectonic setting that can be related to the gradual development of new divergent plate boundaries in the Arctic-North Atlantic and to the onset of convergence of Africa with the southern margin of the European Craton (Ziegler 1990). Compressional tectonics during Coniacian to Campanian times, also referred to as “Subhercynian tectonic events”, caused inversion, particularly of the southern part of the CEBS (Voigt et al. this volume). Further inversion took place in the latest Cretaceous and Early Tertiary. In the LSB, especially the fault systems along the northern and southern basin margins became reactivated and the sedimentary basin fill was thrusted over the adjacent stable Pompeckj Block in the north, and the Münsterland platform in the south (Betz et al. 1987). Probably the LSB played a key role within the CEBS as it might have acted as a barrier zone for extensional deformation triggered from the North Sea rift in the north, and compression derived from the Alpine area in the south (Lohr et al. 2007a). More detailed descriptions on various aspects of the evolution of the LSB have been given by Schott et al. (1967, 1969), Michael (1974, 1979), Kemper (1979), Betz et al. (1987) and Mutterlose (1992). The composition and genesis of the black shale sequences has been discussed in detail by Littke et al. (1998) and Mutterlose et al. (in press). The most up to date review has been compiled by Mutterlose and Bornemann (2000, 2002). Tectonic controls on deposition Differential subsidence controlled by dextral transtension along NW-SE striking faults (Lohr et al. 2007a) resulted

in an asymmetric geometry of the LSB, bound to the north and south by synsedimentary faults with sedimentation rates highest in the centre (Senglaub et al. 2006; Adriasola-Muñoz et al. 2007). Much of this sedimentary sequence (up to 6000 m) was later eroded there. The Berriasian (“Late Cimmerian”) unconformities are evident only along the basin margins. Local tectonics are caused by salt diapirs mainly in the eastern part of the basin and along its western, southern and eastern margins. Particularly in the western Emsland and the Salzgitter area (Fig. 4.3.19) sedimentary patterns vary considerably over less than a kilometre, as recorded by an extreme range of different lithologies. Climate and sea level control The Early Cretaceous Berriasian was characterised by a sea level lowstand (Haq et al. 1987; Hallam 2001) and widespread deposition of non-marine sediments. Coalbearing deltaic and shoreline sediments as well as common dinosaur findings (Iguanodon) and plants in the “Wealden facies” (Bückeberg-Formation) indicate warm humid conditions. A major and long-lasting eustatic sea level rise initiated in the early Valanginian and was succeeded by the early Barremian regression and a shift toward more arid climate (Mutterlose et al. in press) in the late Barremian. The early Aptian transgression was associated with the establishment of new seaways between the Tethys and the Boreal realm via northern France and southern England contemporaneous with a change toward humid climate. Eustatic sea level rise then culminated during the early Turonian with a global sea level maximum of ~100 m above present-day (Miller et al. 2005).

4

Chapter 4.4

Sedimentation during basin inversion T. Voigt · K. Reicherter · H. von Eynatten · R. Littke · S. Voigt · J. Kley

4.4.1 Introduction Basin inversion is a process which is caused by changing intraplate stress orientation and results in compression of the basin fill. The term inversion means the reactivation of faults in a reverse sense (normal faults to thrusts) and the uplift of formerly subsiding areas and viceversa subsidence of former highs. Mostly, “basin inversion” is applied in a much broader sense to whole sedimentary basins, although often only parts are affected. Basin inversion is often caused by compression of the basin, but transpression and strike slip deformation may also generate subsiding areas and uplifts within a basin (Allen and Allen 2005; Kley et al. this volume). In general, basin inversion is reflected in the fragmentation of a basin into subbasins and intrabasinal highs. Basin compartments may be uplifted and eroded, the eroded sediments being redeposited in adjoining depocentres. Clastic deposits of inversion-related sub-basins are therefore often recycled from older sediments. In areas of strong inversion, basement rocks can reach the surface and supply fresh

material to the basin. Basins formed during inversion are of limited size in comparison to epicontinental sag basins or foreland basins. They are often characterised by a narrow elongated shape and a short history of rapid subsidence. This special basin type was first defined by Voigt (1963) who described different examples from central and western Europe. He called a rapidly subsiding, fault-bounded basin situated at margins of stable blocks a “Randtrog” (marginal trough). Detailed investigations (e.g., Stackebrandt 1986; Baldschuhn et al. 1991; Mortimore et al. 1998; Nielsen and Hansen 2000; Kockel 2003; Nielsen et al. 2005; Mazur et al. 2005; Voigt et al. 2004, 2006; Krzywiec 2006b; von Eynatten et al. 2007) allow a more precise view on the origin and evolution of inversion related basins.

4.4.2 Basin Formation Subsidence and accommodation space of inversion related basins are produced by different modes (Fig. 4.4.1).

Figure 4.4.1. Broad and relative shallow marginal troughs form at the fringes of uplifting highs in consequence of crustal thickening due to compression (a) Load of both thickened crust and redeposited sediments is applied to the continental thrust and cause margins to subside (Nielsen and Hansen 2000). Secondary marginal troughs develop after compression due to relaxation and elastic rebound of the lithosphere (b) Depocentres migrate off the uplifted structure (Nielsen et al. 2005)

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and thus leads to crustal thickening. A lithospheric load develops along the inversion axis, although erosion removes sediments above base-level, because light material is eroded and replaced by underlying compacted and thus denser rocks. The primary marginal troughs form as a consequence of isostatic flexural compensation (Fig. 4.4.1a). Subsidence of the marginal trough is enhanced by the load of redeposited sediments. Resulting basins are strongly asymmetric with the deepest axis near the uplifted structure. A secondary marginal trough superimposed on the main inversion-related basin is formed if stress relaxation occurs (Fig. 4.4.1b). Flexural marginal troughs of major inversion zones subside slowly in comparison to fault-bounded basins. Depocentres of the secondary marginal trough are separated from the depocentres of the primary marginal trough.

4.4.2.2 Flexural (thrust load) basins

Figure 4.4.2. Four different modes of the formation of narrow inversion related basins can be discussed, differing in controlling tectonics, sediment architecture and depositional style. Thrust load basins (a) are comparable to foreland basins and develop as a consequence of flexural bending of the lower plate. Subsidence of half ramp basins (b) is provided by rotational thrusting on the back limbs of thrust sheets. Basement folding (c) occurs mainly in the early stages of basin inversion before thrusts start to develop. Sediment accumulation is limited to the synclines. Occurrence of thick salt deposits supports the formation of fault propagation faults (d) by decoupling from basement tectonics. Further compression leads to faulting and allows the salt to pierce the sedimentary succession. Depositional space is provided by salt re-distribution (rim synclines) and leaching of salt (collapse basins)

4.4.2.1 Flexural marginal troughs These structures are commonly developed at the southwestern border of the East European Platform (TTZ and STZ, see Maystrenko et al. this volume) and represent former Jurassic to Early Cretaceous graben structures which became inverted during Late Cretaceous. The basins accompany both sides of shallow anticlines and elongated domes. A model for the formation of broad asymmetric basins flanking large anticlines was proposed by Nielsen and Hansen (2000): Compression and shortening of the lithosphere squeezes the basin fill out of the rift

Basin compression may result in the formation of thrust faults. The load of a thrust sheet causes bending of the foreland plate (Fig. 4.4.2a). Flexural deformation of the plate provides accommodation space for sediments eroded from the hanging wall. Sediment and thrust load cause further subsidence of the thrust load basin, as does the load of sea-water flooding the flexural depression. Deflection of the elastic continental lithosphere leads to uplift of a forebulge. Erosion of the flexural bulge may contribute to the basin fill too if it appears above base level, but more often the bulge shows only reduced sediment thickness if base level remains high. Propagation of thrusts can be detected in a typical coarsening upward pattern in the syntectonic succession. The depocentre is situated near the propagating thrust. Distribution of facies belts reflects the active thrust-sheet; coarsest deposits are derived from the upper plate.

4.4.2.3 Half ramp (piggy back) basins This basin type typically occurs in intraplate thrust belts where several thrust sheets are active at the same time (Fig. 4.4.2b). As most thrusts flatten to a detachment at some depth, shortening causes rotation of the thrust sheets. Accommodation space is provided by downward rotation of the backward limbs. If frontal thrust sheets reach the erosion level, sediment is redeposited on a subsiding half ramp, forming a wedge-like body. If basement thrusts do not reach the surface (blind faults), an asymmetric growing fold develops on top and supplies sediment to the rotating backward limb of the next thrust sheet. The basin fill shows an asymmetric facies distribution with coarser deposits near the uplifting basin margins and fine-grained

Sedimentation during basin inversion Figure 4.4.3. Timing of tectonic activity of a particular structure (simple basement anticline) can be reconstructed from the stacking patterns of sediments. Pre-tectonic deposits show no relation to the growing structure concerning both thickness and facies

deposits in the basin centre. The steep gradient in front of the active thrust is mostly reflected in a narrow belt of aggraded coarse deposits, while the backlimbs are characterised by a gradual transition from nearshore to foreshore deposits. This basin type can be distinguished from flexural basins by the occurrence of a basin set of the same orientation, rotation of a flat basin floor and the lack of a forebulge.

4.4.2.4 Basins due to basement folding Compression may cause broad-scale folding of the lithosphere (e.g., Ziegler et al. 1995). The whole basement and the sedimentary cover form broad shallow anticlines and synclines (Fig. 4.4.2c). While the top of growing anticlines

will be eroded, subsiding synclines may be filled with the eroded material. In the case that growing folds remain below the base level of storm- or current erosion, submarine swells with limited accommodation space develop. Condensed sections of pelagic deposits (chalk) are typical for anticlines in the basin centre. Low sea-level or enhanced uplift of such swells causes non-deposition or erosion and the sedimentary succession becomes incomplete. Eroded or bypassed sediment fills depocentres instead and results in thick successions in the synclines. The load of deposited sediments causes additional subsidence of depocentres. In contrast to basins due to flexure and to piggy-back basins, these synclines should be characterised by a rather smooth, symmetric shape and a central depocentre. Coarse clastics may derive from anticlines on both sides of the basin if appropriate rocks appear at the surface. Figure 4.4.4. Sedimentary pattern around a formerly active structure allows to analyse onset and end of uplift (basement thrust with adjacent thrust-related basin). Syntectonic deposits have a wedge-like geometry with highest thickness close to the active fault. Pre-tectonic and post-tectonic deposits show no relationship to the active structure and to the adjacent marginal trough with respect to both thickness and facies. Note progressive unconformities in the syntectonic basin-fill and basal unconformity of overlapping posttectonic sequences

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Box 4.4.1 Progressive unconformities and timing of deformation events Deposition on growing anticlines leads to reduced thickness of deposits (condensation) because currents prevent the accumulation of sediments and siliziclastic bedload is trapped in the surrounding sinks. The base level falls, either due to enhanced uplift or by falling sea level erosion. An erosive surface is formed which cuts down successively into the previously deposited sediments. Cessation of uplift or rising sea level results in renewed deposition on the erosive surface. If the previously deposited succession was tilted during tectonic activity, an angular unconformity develops. Continuing activity of the thrust then leads to further rotation of both the succession and the unconformity surface. The process of rotation – erosion – deposition may recur several times if base level fluctuates. The main features of these grouped unconformities are a decreasing angle of unconformity with decreasing age and the fact that they disappear basinward in a very complete and thick succession. Such syntectonic unconformities are considered as progressive unconformities. Progressive unconformities very often develop above thrust faults, and have been described from different areas of foreland deformation, such as in the Rocky Mountain foreland and in the Pyrenean fold-and-thrust belt (Spang and Evans 1988; Meigs 1997). Observations of numerous basement thrusts reaching the overlying sedimentary sequences have shown that the focused deformation of the basement fault is mostly distributed to the sedimentary cover, leading to the development of fault propagation folds (Spang and Evans 1988; Suppe and Medwedeff 1990; Allmendinger 1998). If thrusting rates are higher than the rate of accommodation, these folds will be eroded continuously, but if rising base level or decreasing tectonic activity provides accommodation space at the fold margins, rotating unconformities will be produced (Spang and Evans 1988). Stratal geometries across unconformities preserve the tilt of beds established prior to deposition above the unconformity. Thus, a complex but completely redeformable series of angular relationships will be preserved across the unconformities.

4.4.2.5 Rim synclines and collapse basins Rim synclines of salt pillows (primary peripheral sinks) and diapirs (secondary peripheral sinks) may be generated independently from tectonic processes due to the buoyancy contrast to the overlying sequence (see Kukla et al. this volume; Scheck-Wenderoth et al. this volume). Folding of salt-bearing successions or tilting of the basin floor accelerates salt migration in the direction of the maximum hydraulic gradient (Fig. 4.4.2d). Physical properties of salt, in particular low viscosity and its behaviour as a viscoelastic fluid, support decoupling of the basement from the overlying sedimentary cover and allow fault propagation folds to evolve (e.g., Suppe and Medwedeff 1990). Salt is forced to flow from synclines to anticlines by increasing overburden pressure of synkinematic sediments. Hence, salt migration provides accommodation space as well as

causing folding of the sedimentary cover (salt-injected anticlines and salt depleted synclines). Redistribution of salt to the anticlines causes depocentres to migrate towards the anticline and can therefore be differentiated from passive folding. Both propagation of thrusts to the surface and the erosional truncation of overlying rocks expose the salt to water of low salinity. The solution and erosion of salt depends on the water quantity and is thus most effective in marine settings. Dissolution and extrusion of the salt-filled anticline cores cause the anticline to collapse (Fig. 4.4.2d). The overlying sequence bends towards the extrusion centre and allows sediment to accumulate. These basins are of relatively limited size and commonly surround the area where the salt reached the surface.

Sedimentation during basin inversion

All these different modes of basin formation may interact with and amplify each other. In a particular example it might be difficult to decide whether the basin was initially formed by the load of a thrust sheet or rotational tilting on the backlimb of a thrust sheet. Basement folding seems to be a major process for explaining folding and flexural bending of the sedimentary cover in front of major basement thrusts, which break through the growing fold in a later stage and obscure the initial phase (see chapter 4.4.4.2). The existence of thick salt deposits which are involved in basin deformation supports the emergence of fault propagation faults and might prevent basement thrusts from penetrating the basin fill (decoupling). Salt migration also enhances the amplitude of basement folds in the sedimentary cover.

4.4.3 Effects of basin inversion on deposition 4.4.3.1 General remarks Subdivision of a basin into uplifting areas and subsiding troughs has strong influence on sediment distribution even if the uplifts do not reach the surface. Thickness of pelagic and hemipelagic muds will be enhanced in subsiding basins and reduced on swells (Fig. 4.4.3). Uplifts might act as barriers and prevent clastic sediments from accumulating in the sub-basins behind. On the other hand, uplifts contribute to the basin fill if they reach the surface and become eroded (Fig. 4.4.4) Spatial and temporal distribution of thickness and facies of single units reflect activity of growing structures.

Reconstruction of the timing of tectonic activity depends largely on the relationship of sediment bodies to tectonic structures (Fig. 4.4.3 and 4.4.4) Pre-inversion deposits are deformed by the uplifting structure. Their facies and thickness show no relationship with the growing structure. Syn-inversion deposits show reduced thickness on top of swells (condensed sections). Hiatuses (times of non-deposition), erosion surfaces and hardgrounds form if base-level reaches or drops below the sediment surface. Specific shallow-water facies types may develop and interfinger with fine-grained deposits towards the basin. Facies types and their associations in time and space are controlled by uplift rate and size of the active, growing structure. The sub-basins are characterised by complete sections, often with abnormal thickness. Post-inversion sediments are deposited after uplift has ceased and cover both formerly uplifting and subsiding structures. Facies and thickness show no relationship to swells and troughs (Fig. 4.4.4). Nevertheless, secondary effects like crustal relaxation or further salt migration may conceal termination of tectonic activity. No sharp boundary between syn- and post-inversion deposits is developed on structures of slowly decreasing uplift rates and transition from syn- to post-inversion deposits might be difficult to recognise. Active growing structures reaching the surface (Fig. 4.4.4) can be identified by the presence of an erosional surface cutting down into pre-inversion deposits and redeposited clastic material in the surroundings of the growth structure. Uplift rates are commonly high and both pre- and syntectonic deposits are tilted. Progressive unconformities (see Box 4.4.1) are typical for syntectonic deposits, while post-tectonic deposits will cover all tec-

Figure 4.4.5. Pelagic and hemipelagic sediments, which were deposited during Late Cretaceous across Western and Central Europe, derive mainly from calcareous skeletal remains of planktonic organisms. As deposition occurred often below storm wave base, homogeneous chalk deposits of considerable thickness developed

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T. Voigt · K. Reicherter · H. von Eynatten · R. Littke · S. Voigt · J. Kley Figure 4.4.6. Base level changes in the order of several metres had probably no influence on deposition of chalk as is revealed by the preservation of climate-induced cyclicity and nearly complete stratigraphic records. Major base level falls either induced by falling sea-level or by uplift of swells caused slumps and turbidity currents. Hardground formation and redeposition occurred if base-level reached the sediment surface

tonic structures above a major basal unconformity. If uplift rates decrease slowly, late inversion related deposits may be deposited on top of the active structure (above a basal angular unconformity). Deformation of the succession and a regional level higher than deposits of the same age would indicate deposition in the waning stage of uplift. In the following, some aspects of syn-inversion sedimentation with special emphasis on the Central European Basin will be considered.

4.4.3.2 Swells and troughs – condensation and thickness enhancement In the early stages, basin inversion is accompanied by the formation of swells subdividing remaining basin compartments. Initial basin inversion of the Central European Basin coincided with the highest sea-level in earth history (Turonian to Campanian), and thus deposition of pelagic lime-

Sedimentation during basin inversion

Figure 4.4.7. Uplift of intra-basin swells during inversion was followed by massflows which were triggered by steepening of the basin floor. Debrites and slumps deposits of early Coniacian age can be observed at the southern margin of the inverted Lower Saxony Basin (Osning Thrust). They consist of reworked hemipelagic limestones of Turonian age and mark the early stages of basin inversion

Figure 4.4.8. Clastic ironstones of Late Cretaceous age surround the margins of some of the inversion structures. The occurrence of these clasts refers to reworked siderite concretions of Jurassic and Early Cretaceous age (Santonian, Borken, Münsterland Cretaceous Basin)

Figure 4.4.9. Palaeogeographic map of the southern margin of the Central European Basin during the Coniacian to Maastrichtian. Location of inverted areas and post-inversion deposits

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stones prevailed over the basin centre. The process of basin compression started slowly and deformation rates were moderate so that sedimentation occurred both on the swells and in the sub-basins. (Fig. 4.4.5). Most uplifted basin compartments did not reach the erosion level which is best confirmed by the storm wave base for unconsolidated mud in open marine environments. The storm wave base depends strongly on the configuration of the sea and climatic conditions. It fluctuates between 100 and 200 m both in recent epeiric seas and at open-ocean shelves (e.g., Johnson and Baldwin 1996). Erosion of hard rocks requires much higher energy and is mostly limited to the upper shore face because the depth of surf action does not exceed a few metres in most marginal seas (e.g., Reading and Collinson 1996). Swells below stormwave base in open marine environments get their sediments from different sources (Fig. 4.4.5). Remains of planktonic organisms, especially calcareous shells, provide the majority of the sediment supply (pelagic rain). Coccoliths, foraminifers and calcispheres are the main components of the pelagic rain during the Late Cretaceous (Wilmsen 2003). Additionally, clay minerals are deposited from suspended river input. Clay content of pelagic deposits depends mainly on the distance to the ancient coast, because the salinity of marine water supports aggregation of larger clay flakes (flocculation). Coarser clastic deposits like sand and silt which are predominantly transported as bed load are trapped in the surrounding areas of higher subsidence and do not reach the top of swells. Carbonate content of swell deposits is therefore higher and thickness is lower in comparison to subsiding sub-basins. Quartz silt, nevertheless, is the major component of aeolian dust which may be transported over high distances and contributes to the sediment budget. The influence of aeo-

lian dust, however, was small in the Cretaceous greenhouse world with its limited desert areas. If the top of swells reach the level of current action (mostly storm wave base) either due to falling sea-level or enhanced uplift rates, sustained accumulation of sediments is impossible (Fig. 4.4.6). Erosion cuts through unconsolidated mud down to lithified sediments that are redeposited in surrounding troughs (Hancock 1989). These “hardgrounds” may form a stable sea bottom over several thousand or even million years. Rising sea-level again allows the deposition of pelagic mud which covers the hardground. Storm wave action re-suspends unconsolidated chalk and the resulting sediment clouds may transform into turbidity currents, move basinward over large distances and settle as turbidites. Most of these turbidites must be considered in fact as distal storm beds because storm action was probably the trigger for resedimentation. Typical structures of turbidites and chalk storm deposits are normally eradicated by burrowing marine animals bioturbation). Ekdale and Bromley (1991), however, were able to show that tiering patterns of trace fossil communities provide evidence of erosion, hiatuses and event-sedimentation even in such homogeneous sediments like chalk. A second important process for sediment redistribution is slope failure. Steepening of the sea-bottom by swell uplift or trough subsidence causes slope instability leading to gravitational massflows which accumulate in the troughs. Slumps, slides and debris flows of considerable thickness may develop if unconsolidated sediments start to move (e.g., Mortimore et al. 1998). Sheet-like bodies of mass flows have been frequently observed near active thrusts (for example in the early Coniacian and late Campanian Figure 4.4.10. Pre-Tertiary subcrop of the Münsterland Basin south of the Lower Saxony Basin at the southern rim of the CEBS

Sedimentation during basin inversion

of the Münsterland Basin, Voigt 1963) and are thus good indicators of syntectonic deposition. (Fig. 4.4.7). Turbidity currents related to such mass flows also contribute to sediment dispersion but may be difficult to distinguish from distal tempestites. A special type of swell sediment in the described example of the Central European Basin is represented by shallow water carbonates which were deposited around slowly uplifting highs or on top of inversion structures after inversion had stopped. In particular, the Campanian deposits on top of the inverted Lower Saxony Basin (Stemwede Formation), isolated swells of the Münsterland Basin (Burgsteinfurt Formation), and the Danian chalk of Denmark and South Sweden are composed of coarse-grained pure limestones. Bivalves, red algae, echinoids and bryozoans contribute to bioclastic limestones which formed in isolated shallow water environments between fair-weather wave- and storm wave base.

4.4.3.3 Clastic deposition – the result of uplift and erosion If the uplift rate of inverted structures is higher than the rate of sea-level rise, uprising highs may grow above sea level and will then be eroded progressively. The weathered and eroded material is transported to the surrounding troughs, with sediment composition largely depending on the bedrock composition of the uplifting high. Some of the inversion related basins like the Münsterland Basin are nearly completely filled with mudstones and thus reflect the very thick Early Cretaceous pelitic successions of the inverting Lower Saxony Basin. Since most of the inversion structures originate from former rifts, a thick cover of sedimentary rocks has to be removed before basement rocks are involved in erosion. If inversion slowly exhumes older coarsegrained clastic deposits (e.g., Early Triassic Buntsandstein), sediment recycling leads to deposition of mature quartz-rich sandstones. Deeper exhumation leads to exposure of basement rocks that may be identified by typical heavy minerals. In the case of the Central European Basin, only a few basement uplifts reached the surface during inversion. All of them are situated at the southern margin of the basin where thickness of sedimentary cover is moderate and where the maximum exhumation rates occurred (Ibbenbüren High, Harz Mountains, Lausitz High, Thuringian Forest, Kyffhäuser). One of these structures, the Harz Mountains will be described in more detail in chapter 4.4.4.2. The amount of sediments in the new sub-basins depends on the size of the uplifted areas. Large shallow anticlines rimmed by narrow marginal troughs produce more sediment than accomodation space (overfilled basins). Additi-

onal sediment will be spread over the depocentres and deposited in other slowly subsiding areas. In this case, facies distribution is independent of the position of the depocentre. The opposite case, small uplift structures surrounded by subbasins which are able to catch more sediments than are provided by the source area, leads to underfilled basins which become deeper during inversion. This seems to be the case in the eastern part of the Münsterland Basin, which is described briefly in chapter 4.4.4.1. Coastal facies belts are narrow and strictly controlled by the position of the uplifting structure. While relatively coarse-grained clastic material will be trapped in the surrounding basins, suspended and dissolved matter spreads over larger areas. Coarse material, which is delivered by relatively short rivers and is thus of relatively low compositional and textural maturity, can be deposited in immediate vicinity of the river mouths if active thrusting occurs (e.g., Ulicný ˇ 2001). Examples of such successions can be found at the inverted southern margin of the Central European Basin. Thick sequences of conglomerates and coarse sands were deposited around uplifting highs at the margins of the Bohemian Massif (Saxonian-Bohemian Cretaceous Basin, Innersudetic Basin, Regensburg Basin, Westsudetic Basin). These clastics fill the depocentres in front of thrusts and pass over short distances into marly offshore successions. Narrow facies belts of coastal sands and the sudden transition to bioturbated low energy offshore sediments indicate steep gradients of the basin floor of an under-filled basin. Sediments are distributed by currents and storms if open marine conditions at high energy coasts prevail. If moderate thrusting rates prevail and the depocentres are at some distance from the shore, well-rounded quartz sands of high textural and compositional maturity originate from recurrent wave action. Protected coasts with beach-barrier sequences and coastal plain deposits (including fluvial sediments and coals) provide evidence of low basin-gradients and high sediment supply from surrounding inversion areas. Coastal plain deposits spread towards basin centre if accommodation space decreases. Progradation is forced by falling sea level, decreasing subsidence or enhanced sediment supply.

4.4.3.4 Ironstones and phosphorites around inversion structures and diapirs A special sediment-type of former economic importance occurs near Late Cretaceous inversion structures of the Central European Basin. Ironstones of minable thickness are limited to intra-basinal highs with moderate or decreasing uplift rates. They form intercalations within

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sandy and marly successions of the Santonian to Campanian and consist predominantly of angular limonite clasts of cm-size (Fig. 4.4.8). Most of the ironstones were probably reworked from older clastic ironstones of the Early Cretaceous, but also lateritic weathering of Jurassic and Early Cretaceous iron-rich sediments (in particular siderite nodules) contributed to the iron ore deposits (Kolbe 1962). Ironstone formation depends on mobilisation of iron under tropical humid climate and subsequent precipitation in the soils or in the sea. The pebbles of dense limonite were concentrated at the shores of the Late Cretaceous sea around active highs and finally deposited in the surrounding rim synclines (Niebuhr et al. 2000). Additionally occurring ooids point to primary precipitation of iron hydroxides in sea-water and shallow water reworking. Preferred occurrences of such ironstones are the margins of larger inversion structures with relatively low subsidence rates, for example, the western margin of the inverted Subhercynian Block (Lengede-Broistedt), the Altmark Basin north of the same unit and the western Münsterland Basin. Phosphorites occur within the Coniacian Halberstadt formation at the northern border of the Subhercynian Cretaceous Basin. They derived from transgressive Cenomanian deposits (Voigt 1929) and were reworked in the earliest stages of inversion in the early Coniacian. The clastic phosphorites are of brownish to yellowish colour and occur as pebbles of phosphorite-impregnated quartz sands,

pure phosphorite nodules or even as phosphorised fossils, like ammonites or bivalves. Primarily, most of these phosphorites formed as concretions within condensed successions of Albian and early Cenomanian glauconite-rich deposits.

4.4.4 Sedimentation during inversion in the Central European Basin NE-SW compression of the Central European Basin lead to large scale folding of the basement and the overlying basin fill during Late Cretaceous. NW-SE striking normal faults which developed during Late Jurassic to Early Cretaceous basin differentiation were reactivated to reverse faults. As deformation was dominated by shortening and minimum stress was in a vertical direction, large areas were uplifted above base level and subjected to erosion. Principal deformation was focused to the north-eastern and the south-western margins of the CEBS. The STZ and the TTZ (see Fig. 2.5, Maystrenko et al. this volume) with about 2000-3000 m of uplift and about 1500 m of syn-inversion deposits in the surrounding marginal troughs (up to 2000 m in rim synclines of active diapirs) were investigated recently by Nielsen et al. (2005) and Krzywiec (2006b), pointing to an overall similar inversion history along strike modified by salt tectonics and occasional thrust faults. The south-west-

Sedimentation during basin inversion

Figure 4.4.12. Thickness of eroded Mesozoic and Zechstein deposits in the western Lower Saxony Basin (after Senglaub et al. 2005)

ern margin shows a very complex fault pattern accompanied by very different subsidence and exhumation rates. Timing and amount of inversion is partly difficult to assess due to widespread Cenozoic erosion. Recent studies focused on the seimic lines across the inverted Lower Saxony and Altmark Basins (Mazur et al. 2005), analysis of the uplift and subsidence history of inverted basins (Petmecky et al. 1999; Senglaub et al. 2005) and integrated investigations of basement uplifts and related basins (Voigt et al. 2004, 2006; Franzke et al. 2007; von Eynatten et al. 2007) and will be referred in the following. The palaeogeographic map of the southern margin of the Central European Basin during the Coniacian to Maastrichtian (Fig. 4.4.9) reflects an en echelon pattern of broad anticlines which formed extended islands in the Late Cretaceous chalk sea and were surrounded by a broad belt of marly deposits. The former Central Netherland Basin, the inverted Lower Saxony Basin (WestfalenLippe-swell), the Prignitz-Lausitz uplift and the Subhercynian Block were 100 to 300 km long and 50 km wide. The large scale swells and anticlines are subdivided by smaller structures, 20-150 km long and some 50 km wide, bounded by thrust faults (e.g., Flechtingen High and Harz Mountains). In contrast to the flexural deformed largescale uplifts with moderate uplift rates and limited exhumation, these thrusts led to rapid transport of basement blocks to the surface. Most of these highs represent areas of Late Jurassic and Early Cretaceous sedimentation and were partly stripped of the sedimentary cover. Seismic and borehole data show enhanced thickness of Jurassic and Early Cretaceous sediments in some of the Late Cretaceous swell areas if the amount of inversion was limited. In most cases, however,

erosion cuts down to the Triassic and the initial sediment thickness can only be concluded indirectly. This is possible by the reconstruction of the thermal history of the preserved succession (see Littke et al. this volume) or by investigation of the detrital sediment components of the adjacent syn-inversion basins (see Box 4.1.2). The boundaries of the inversion structures are different: shallow to steeply dipping flexures are most typical, but some of the structures are bounded by major thrusts (Fig. 4.4.9). Uplift of thrust sheets and anticlines transported sedimentary rocks to the surface where they underwent weathering and erosion. Eroded rocks were deposited in the vicinity of the uplifted areas, as is reflected in a broad facies belt of marls (Fig. 4.4.9). The coarse-grained facies belt of coastal sands, however is very narrow corresponding to the fact that the basin axis with highest subsidence was situated immediately at the front of active thrust sheets. Sands are restricted to the western Münsterland Basin, to the Subhercynian Basin and to the Altmark Basin, and mainly reflect erosion of marginal sandstone units of the Early Triassic and Early Cretaceous. Reworked bed load sediments were trapped in the depocentres and only suspended material (clay and silt) could spread across the Central European Basin. Clastic sediments were restricted to the southern margin of the basin although similar inverted structures and uplifts also existed in the central basin. The reason is a decreasing shortening rate and a general plunge of the basin floor towards the north, so that uplifted thrust sheets remained below sea-level. Only gradual thickness changes of chalk deposits or cyclic limestone-marl alternations of the transitional facies belt reflect the existence of growing and subsiding areas (Niebuhr 2005). At the northern border of the basin, the Sorgenfrei-Tornquist Zone inverted in a similar way to the large structures in the south, but the

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Figure 4.4.13. The 90 km long Subhercynian Cretaceous Basin is clearly related to the Harznordrand Thrust. Syntectonic deposits reach more than 2500 m thickness in the central basin and pass into hemipelagic deposits towards the Northwest. The whole area is influenced by syn- and post-sedimentary compressional structures

amount of inversion was less than at the southern margin. Extended marginal troughs formed due to flexural loading and were filled with thick chalk successions and marly deposits resulting from redeposition of basinal mudstones of Early Cretaceous and Jurassic age. As sea-level was exceptionally high during the Late Cretaceous (the highest in earth history), the majority of sediments were deposited in a marine environment. Several basins at the southern margin of the Central European Basin reflect fully established marine facies belts starting with coastal sands and conglomerates grading into nearshore sands and bioturbated foreshore mud. Evidence of storms and tidal transport in some nearshore units point to an open marine environment with broad connections to the open ocean. The composition of redeposited sediment is mainly determined by the composition of eroded bedrock. A high proportion of the previous basin fills was composed of Mesozoic pelitic rocks and, therefore, redeposited sediments largely consist of clay and silt. Mixtures of biogenic material and marls prevail (marls of the Emscher Formation). Evaporites were dissolved during weathering. Likewise, Triassic and Jurassic limestones occur

only as pebbles in immediate vicinity of the active faults. Thick sandstone sequences derived mainly from reworked Buntsandstein and from frequently recycled Early Cretaceous sandstones. The deposits in the surroundings of inversion structures represent an “unroofing sequence”, that means, the material of the sedimentary cover occurs in reverse stratigraphic order in the sediment of the newly created basins (Box 4.4.2). In the following section, two examples of inversion-related basins are described, differing with respect to their position within the basin system, tectonic setting and depositional style of basin-fill.

4.4.4.1 Basins related to inverted grabenstructures – the Münsterland Basin The Osning Thrust, more than 100 km long, is one of the most prominent tectonic structures at the southern margin of the Central European Basin (see Kley et al. this volume). It represents the reactivated main boundary fault of the Lower Saxony Basin and separates the adjoining Münsterland Basin in the south, where more than 2000 m of syntectonic sediments were deposited (Hiss 2000). This

Sedimentation during basin inversion

Figure 4.4.14. A long-strike cross-section of the Subhercynian Basin and the adjacent basement uplift of the Harz shows a symmetric shape of the basin and a considerable offset of the basement surface at the thrust which might surpass 10 km. The trace of the eroded basement surface above the Harz Mountains was reconstructed by extrapolation from adjacent basins and by fission track data

basin is a well-investigated example of a marginal trough developed in front of a fault-bounded inversion structure and will therefore be explained here in some detail. The Münsterland Basin was established on the Rhenish Massif, which remained a stable block during most of the Mesozoic and acted as a source area in Late Jurassic and Early Cretaceous (Littke et al. 1994; Büker et al. 1995; Karg et al. 2005). A global sea-level rise caused Albian and Cenomanian deposits to transgress on the Rhenish Massif. They overlay an Early Cretaceous peneplain of folded Late Carboniferous units. The age of peneplanation is proven by remains of Early Cretaceous sediments in caves and incised valleys below the transgressive surface. Facies and thickness distribution of the Cenomanian are

independent of the later basin contours - they belong to the pre-inversion succession (Wilmsen 2003). Differentiated basin subsidence started slowly in the Turonian which is reflected in the enhanced thickness of pelagic limestone deposits along the basin axis. Strong subsidence of the basin axis in front of the reactivated southern margin of the Lower Saxony Basin formed a large asymmetric syncline with a pronounced depocentre in the northwest corner of the basin (Fig. 4.4.10). This area of maximum thickness is situated directly opposite to the area of maximum uplift of the Lower Saxony Basin. Deposits of the Münsterland Basin derived from two sources: the inverted Central Netherland Basin in the west served as the main source of sand and the inverted Lower Saxony Basin north of the Münsterland Basin acted as a source area for mud.

Figure 4.4.15. Stratigraphy of the Subhercynian Cretaceous Basin (modified from Voigt et al. 2006). Cenomanian to Turonian deposits are dominated by pelagic limestones without significant siliciclastic input. The majority of Coniacian to Campanian deposits is made of sandstones and marls

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Figure 4.4.16. Facies distribution within the Subhercynian Cretaceous Basin in space and time reflects the transition from pre-inversion deposits to syn-inversion sedimentation (modified from Voigt et al. 2006; for key see Fig. 4.4.18). Earliest differentiation of the broad Cenomanian shelf is observed in the Turonian by enhanced thickness. Lower and middle Coniacian deposits were shed from the Northeast, activity at the Harznordrand Thrust is not proven

Late Jurassic and Early Cretaceous sediments were deposited in great thickness in the Lower Saxony Basin, but subsequently eroded to a large extent. This erosion occurred mainly in the Late Cretaceous due to basin inversion. The timing of the earliest erosion/inversion has been dated based on mass flows as early Coniacian (Voigt and Koch 1977). Furthermore, fission track ages revealed a strong cooling of the rocks of the Lower Saxony Basin in Coniacian to Santonian times (Petmecky et al. 1999; Senglaub et al. 2005). During this inversion, the basin fill was uplifted and thrusted onto the Rhenish Massiv in the south (Münsterland Basin) and on the Pompeckj swell in the north. The total thickness of eroded sediments was analysed using numerical basin modelling calibrated by maturity and fission track data (Littke et al. this volume). According to this analysis, up to 6000 m were eroded in the basin centre (Fig. 4.4.11), i.e., in the area of Bramsche close to Osnabrück. Because of this, dense and highly mature rocks occur at the Earth’s surface. The centre of the inversion with the greatest thickness of eroded rocks coincides with the long-known maturity and gravity anomaly of the Lower Saxony Basin.

Major faults running along strike (E-W or ESE-WNW) often limit blocks of enhanced/reduced erosional thickness, proving that they were active during the inversion phase. The basic data for individual wells and profiles have been published in Leischner et al. (1993), Petmecky et al. (1999), Senglaub et al. (2005) and Adriasola-Muñoz et al. (2007). Exceptionally high eroded thicknesses were observed in the areas of the Piesberg (7300 m, near Osnabrück) and Ibbenbüren (4400 m), where Carboniferous horst blocks reach the Earth’s surface (Fig. 4.4.12). Missing Zechstein salt in this area may have caused the enhanced inversion and erosion in this region (see Drozdzewski 1988; Ziegler 1990). The Münsterland Basin became successively filled with a mixture of clastic and carbonatic deposits. Thickness distribution reflects flexural bending of the foreland plate. Mass flows developed in the late Turonian and early Coniacian succession of the recently upturned northern basin margin (Voigt 1963). Synsedimentary breccias, debris flows of semi-lithified carbonates and slump folds indicate tilting of the basin floor. The following monotonous

Sedimentation during basin inversion

Figure 4.4.17. Syn-inversion deposits of the middle Coniacian show a progradational pattern from the northeastern source area into the deeply subsided but underfilled Subhercynian Cretaceous Basin (modified from Voigt et al. 2006). The basement block of the Harz Mountains, forming the most prominent structure today, had no direct influence on sedimentation in the Coniacian

marls of Coniacian and Santonian time reach great thickness near the thrust (approximately 1500 m represent only 5 Ma) and pinch out towards the south (Hiss 2000). These foreshore mudstones are strongly bioturbated and only rarely exhibit indications of current reworking. They were deposited below the stormwave base in an open marine environment. The basin was under-filled and sediment supply more or less kept pace with subsidence. Sandy deposits of the Coniacian and Santonian are rarely preserved, pointing to subsequent erosion of corresponding facies belts. Surprisingly, sands are limited to the west, related to the eastern margin of the inverting Central Netherland Basin and to the south, where the passive basin margin was situated. Adjacent to the strongly uplifting Lower Saxony Basin, sands occur very sparsely, dispersed in the thick marl-sequences of the Emscher Formation. This abnormal facies distribution can be explained by the exhumed fill of the inverted Lower Saxony Basin: thick muddy sequences were deposited in the graben-structure in the Early Cretaceous, sands were restricted to a narrow belt near the Rhenish Massif (Osning-Sandstein, Mutterlose 2000), where they were preserved in the marginal flexure between both basins. The predominance of pelites over sandstones in the exhumed basin is expressed in the subsequent basin-fill.

Campanian deposits are much more diverse than the Coniacian through Santonian succession (Hiss 2000). Sandy deposits spread from the west across the basin pointing to progradation of sediments during times of decreasing accommodation space. This could be caused either by tectonics (decreasing subsidence of the basin floor or enhanced uplift of the source area) or by falling sea-level. Facies belts in the west, where nearshore sands were deposited and even conglomerates are preserved grade eastward into storm controlled shoreface deposits. To the east, near the Osning Thrust, similar storm-induced deposits occur but their composition is different from the west. High bioproductivity and higher input of mud instead of sand is observed compared to the older deposits of Santonian to Campanian. Shallowing during Campanian is indicated by the deposition of frequent calcareous stormbeds in marl sequences (Beckum beds of the Ahlten Formation and the bioclastic limestones of the Baumberge Formation). Internal subdivision of the basin by troughs and swells corresponds to further shortening by folding. Swells are characterised by hiatuses and in general thinner successions compared to the troughs. Shallow water bioclastic carbonates predominate while troughs contain bioclastic beds reworked from the swells within thick marl sequences

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Figure 4.4.18. The exhumation of the Harz is reflected in changing facies distribution of the Subhercynian Basin during Santonian and early Campanian (modified from Voigt et al. 2006). Coarse-grained shallow water deposits derive from the sedimentary cover of the uplifting anticline and, later, from the Hercynian basement of the Harz. Several unconformities near the thrust fault and at the northern margin of the basin are developed

(Burgsteinfurt Formation). A relatively steep submarine relief is proven by the occurrence of soft sediment deformation in the Beckum beds. The sedimentary record of the Münsterland Basin ends in the late Campanian but subsequent deformation by folding and faulting could indicate persisting compression and inversion. Transgressive late Campanian and Maastrichtian shallow water limestones of the Damme syncline on top of the inverted succession of the Lower Saxony Basin mark the end of major uplift.

4.4.4.2 Basin evolution in front of a basement thrust: The Harz example The Harznordrand Thrust, one of the major NW-SE striking fault zones of the Central European Basin, divides as a large basement block of the Harz Mountains (Palaeozoic low grade metamorphics and Late Carboniferous igneous

rocks) from the adjacent Subhercynian Cretaceous Basin which is filled with Late Cretaceous deposits (Fig. 4.4.13). Surface outcrops, geophysical and borehole data prove a relatively steep-dipping, northward-directed thrust at the southeastern basin margin. Thermochronological data obtained from the exposed basement point to pronounced uplift of at least 5000 m in a confined period from 8580 Ma (Santonian to Campanian; Thomson et al. 1997). According to borehole data and seismic lines across the thrust, the basement surface (base Zechstein) is at a depth of more than 4000 m in front of the central thrust (Franzke et al. 2007; Voigt et al. 2004). The adjacent Subhercynian Cretaceous Basin represents one of the key structures of Late Cretaceous intraplate deformation of Central Europe. A well-exposed 2500 m thick basin-fill of syntectonic deposits reflects the tectonic history of the area and allows a precise reconstruction of timing and amount of uplift of the surrounding source areas, especially the exhumation of the Harz. The formation of the Subhercynian Cretaceous Basin started in the Turonian

Sedimentation during basin inversion

Box 4.4.2 Exhumation and Unroofing Sequences Exhumation means relative uplift with respect to the earth surface (England and Molnar 1990). The former overburden of the exhumed rock body may be removed either by faulting (tectonic denudation) or by erosion (erosive denudation), or combinations of both. Erosive exhumation and denudation of structural blocks usually leads to synchronous deposition of the eroded material in related sedimentary basins, e.g., foreland basins. Continuous erosion of an exhuming rock pile implies progressive down-cutting into deeper seated rocks, for example, progressively older sedimentary rocks or higher-grade metamorphic rocks. In the adjacent sedimentary basin the succession of the eroded rocks will be inverted (see Fig. 1): near-surface rocks in the source area are eroded first and deposited relatively deep in the basin (blue), whereas the material from deeper seated and later eroded rocks is deposited in the youngest sediments (green). Such sedimentary successions are typically called unroofing sequences (normal unroofing sequences, cf. Colombo 1994). The best way to recognise exhumed blocks is the identification of specific source lithologies in stratigraphically well-constrained sediments. If conglomerates are available, source rock identification is easy. Unfortunately, pebbles of appreciate sizes are rare and are usually restricted to the basin margins and to source lithologies that produce pebbles and boulders hard enough to resist abrasion in high energetic environments. In particular, high relief, steep shores, strong storms and the predominance of resistant lithologies provide good possibilities for the deposition of pebble-sized sediments in the basin. Sandstones are much more common in sedimentary basins, but source rock information is diluted due to disaggregation into sand grains and selective alteration. Many techniques, however, allow the initial source rocks to be reconstructed from sediment bulk composition and single-grain characteristics (see Box 4.1.2). Among the latter, detrital thermochronology is eminently valuable (e.g., Bernet and Spiegel 2004). These techniques allow for constraining both the time and the rates of cooling in the source area for a certain time slice in the geological record (i.e., a stratigraphic interval). Assuming a reasonable geothermal gradient cooling data may be converted to exhumation rates.

and continued into the Campanian (about 10 Ma), as suggested by a distinctive thickness trend indicating a narrow NW-SE orientated subsidence axis in front of the thrusted basement block of the Harz (Fig. 4.4.14). The up to 1500 m thick succession of Coniacian to Santonian deposits represents the short timespan of approximately 5 Ma pointing to rapid subsidence and deposition. Inversion continued presumably at a comparable level through the Campanian, but the preserved sedimentary record ends already with early Campanian deposits. Subsidence and uplift rates are similar to rates observed elsewhere, especially in the Münsterland Basin and the inverted Lower Saxony Basin but are also comparable to estimated rates at the TTZ. The thickness of sediments is highest near the thrust thus forming an asymmetric syncline with a steep, partly inverted limb close to the thrust (Fig. 4.4.4, 4.4.13). The basin history is reflected within several shallow marine

depositional units each characterised by distinct facies associations (Fig. 4.4.15). While Cenomanian and Turonian deposits are dominated by pelagic limestones without major siliciclastic input, the majority of Coniacian to Campanian deposits are dominated by sandy and marly deposits, which were derived from adjacent source areas. Coniacian storm-dominated successions of an under-filled phase of basin development and Santonian deposits of a low-energy delta plain to tidal-flat environment consist of sands that were predominantly shed from the northeast and east (Fig. 4.4.16). In the early stages of inversion, during Turonian and Coniacian, the later Harz Mountains had only minor influence on deposition and acted as a swell area (Fig. 4.4.17). It was not before middle Santonian that facies belts were modified by the activity of the Harznordrand Thrust as indicated by the contribution of freshly exhumed Mesozoic rocks from the former sedimentary cover of the Harz to the sediment deposited

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in the basin (Fig. 4.4.18). During the early Campanian, the eastern source area lost its influence and the exhumed basement south of the Harznordrand Thrust (Harz Mountains) acted as a source area (distal rocky shore deposits). The observed changes in facies and the varying activity of different source areas reflect the interaction of eustacy and tectonics. Syntectonic deposits are preserved in the immediate vicinity of the active fault and reflect the evolution of a continuously growing anticline (fault-propagation fold) above the Harznordrand Thrust that was eroded and acted as a source area during low sea-level. Recurrent periods of high sea-level led to subsequent deposition on the uplifting swell. As the underlying sequence was tilted during tectonic activity, progressive unconformities on the rotated sediments in front of the thrust could form. As is shown in Box 4.4.1, these deposits allow the very precise reconstruction of fault activity in space and time. Close to the fault, at least four angular unconformities were demonstrated (e.g., Voigt 1929), spanning a period of less than 5 Ma (middle Santonian to early Campanian). The overlying transgressive deposits above the unconformities near the Harznordrand Thrust reflect a situation where the increase in accommodation space kept pace with or exceeded the rate of thrust movement. This can be caused either by decreasing thrusting rates or by rising sea-level. In times of decreasing accommodation potential, the uplifting units in front of the thrust zone were eroded and redeposited north of the fault. The preservation of several unconformities very close to the active thrust demonstrates that the thrust developed rapidly from its blind stage (fault propagation fold) to a nearly stationary active thrust, which caused only minor rotation of the footwall. In the case of the described example, the ages of angular unconformities correlate well with the regional and even global pattern of transgressions (Niebuhr et al. 2000). The composition of the sediments reflects the stepwise exhumation, erosion and redeposition of bedrock. The sedimentary sequence contains the eroded section of the hanging wall in an inverted pattern. First, Cretaceous soft sediments were eroded, followed by Jurassic and Triassic, each leaving behind a particular trace in the redeposited succession. After erosion of the Permian limestones and evaporites (Zechstein), basement rocks reached the erosion level and were delivered into the adjacent dump, i.e., the oldest components are redeposited in the stratigraphically youngest unit of the basin fill (“unroofing sequence”, see Box 4.4.2). Provenance analysis of the clastic basin fill, integrated with facies analysis and structural data as outlined above, allows for reconstructing mechanisms and timing of the exhumation of the Harz (Fig. 4.4.19). From latest Turoni-

an up to the early Santonian there is no evidence of detrital input from the nearby Harz area. Sediments are mature quartz arenites derived from intense weathering and recycling of Permian to Mesozoic sedimentary rocks exposed to the east of the Subhercynian Basin (Fig. 4.4.19a). In the middle Santonian the situation changed fundamentally because freshly exhumed and eroded Mesozoic sedimentary cover rocks of the Harz were delivered into the basin (Fig. 4.4.19b,c,d). Feldspar and Mesozoic lithoclasts reflect erosion of Triassic and, in places, Jurassic and Early Cretaceous to Turonian strata. Apatite and garnet in heavy mineral spectra are derived from largely unweathered Early Triassic Buntsandstein as indicated by the analogy of apatite and garnet chemistry (von Eynatten et al. 2007). In the early Campanian, Palaeozoic lithoclasts and the reoccurrence of mostly ultra-stable heavy mineral associations indicate erosion cutting down into the basement of the Harz (Fig. 4.4.19e). A contemporaneous strong decrease in feldspar, garnet and apatite suggests the almost complete removal of the 2-3 km thick Mesozoic cover of the Harz within only 2-3 Ma (von Eynatten et al. 2007). This translates into an exhumation rate of approximately 1mm/a consistent with apatite fission track data from granitoid rocks of the Harz Mountains.

4.4.5 The North German Basin during the Tertiary Timing of inversion, especially the end of inversion of the CEBS, is still under discussion. High quality reflectionseismic lines across the inverted Polish Trough (Krzywiec 2006b) and the Danish Basin (Nielsen et al. 2005) give evidence that inversion occurred largely during Late Cretaceous as in the described examples above, while some authors believe that compression persisted until the Middle Palaeogene (Ziegler 1990; Worum and Michon 2005; Krzywiec 2006b). At least at the southern margin, timing is good constrained by fission track data and distribution of regional unconformities to have been occurred between Turonian and Maastrichtian. Palaeogene sediments as the Danian chalk and sandy deposits of the Zahna Formation in more marginal settings display still the NW-SE orientation of inverted structures, but were deposited in shallow, distal subbasins of domal uplifts. Nielsen et al. (2005) therefore explained that these deposits are not related to compression but to plate-wide relaxation of the in-plane tectonic stress (secondary marginal troughs; Fig. 4.4.1). The Tertiary period in the North German Basin is characterised by the post-inversion consolidation of the crust, partially under extension, and later by the Alpine orogenic phase from the Neogene onwards. The basin-wide inversion due to far-field collisional influence from the Pyrenean

Sedimentation during basin inversion Figure 4.4.19. Illustration of middle Coniacian to early Campanian tectono-sedimentary evolution of the Subhercynian Cretaceous Basin (SCB) (modified after von Eynatten et al. 2007). The integrated investigation of sedimentary facies and provenance allows for precisely reconstructing the variations in sediment dispersal related to the exhumation of the Harz Mountains. Not that in early stages of compression and uplift (A and B) flexuring of the Palaeozoic basement led to the formation of an anticline before onset of thrusting of Palaeozoic rocks onto Mesozoic units in latest Santonian (D and E). In late Santonian (C) the rate of sea-level rise exceeded the uplift rate of the Harz leading to the formation of an unconformity between Heidelberg and Heimburg formations (Voigt et al. 2004). Major provenance indicators are ZTR (zircon+tourmalin e+rutile), grt (garnet), ap (apatite), Qt (total quartz), F (feldspar), and Mesozoic (Lmes) and Palaeozoic lithoclasts (Lpal)

orogeny and the Cretaceous Alpine orogenic phase is followed by a phase of extension, which is not represented in all parts of the basin. The North East German subbasin shows no evidence for extension during that period. During an Early Palaeocene extension phase the area of the North German Basin (Fig. 4.4.20a) subsided and was

flooded, where marine sediments were deposited far into the region of Poland. Scandinavia is uplifted since the Palaeocene and serves since then as a sediment source area for the North German Basin (Nielsen et al. 2002). The initiation of the uplift in Scandinavia is linked to the emplacement of the Iceland

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Figure 4.4.20. Palaeogeographic evolution of the depocenters in the North German and Polish Basins during the Palaeogene from Reicherter et al. (2008; after Vinkena 1988; Stankowski 1996; Peryt and Piwocki 2004). Note the presence of a NW-SE trending subsidence center (North German Basin) and a topographic barrier (Mid-Polish Swell) until the Late Eocene. A significant change occurred during the Oligocene into an East-West elongated depocenter

plume about 53 Ma ago which led in consequence to heating and lithospheric delamination (Nielsen et al. 2002). However, variation in sediment supply and distribution are influenced by climatic variations and a general falling eustatic sea level in the Late Palaeocene (Fig. 4.4.20b). During the Eocene a transgression due to significant sea level rise flooded most of the North German Basin (Fig. 4.4.20c), and created a seaway connection with the Dnepr-Donets basin. In Germany major subsidence took place

in the Hamburg area, here also the influence of salt diapirism (e.g., the Glückstadt Graben) is observed. In eastern Germany along the basin margin, widespread lignite deposits formed in the area of Halle and Leipzig. Middle to Late Eocene deposits and biofacies suggest that during this time the Mid-Polish Swell still existed as an elevated zone (Fig. 4.4.20c). In particular, the dominance of the Mediterranean and Boreal faunal elements documented to the SW and NE of the Mid-Polish Swell indicates that

Sedimentation during basin inversion

Figure 4.4.21. Amplitude of uplift or subsidence since beginning of the Rupelian (Early Oligocene; approx. 34 Ma ago) in the North German Basin; from Stackebrandt (2004). Note the NW-SE trending center of subsidence, with intrabasinal NNE-SSW trending faults, and a general deepening of the basin towards NW

there was a topographic barrier in between the two zones. Periodic mixing of the two faunas suggests a relatively low relief of the swell. The end of the Eocene is characterised by uplift along the southern and northern margins of the North German Basin, accompanied by a sea level drop. A sudden but high-magnitude sea level rise in the Early Oligocene flooded the basin again, now it reaches its main extension linking the North Sea with the Upper Rhine Graben via the Hessian Depression. The Rupelian deposits (Rupelian clay, “Rupelton”) represent a classical marker horizon to deduce uplift/subsidence rates, i.e., vertical crustal movements since Oligocene times (Fig. 4.4.21). Marine sediments of the Early Oligocene are preserved in the Harz mountains close to Elbingerode at an altitude of 512 m above sea level, whereas the same sediments are found approximately 1000 m below sea level close to Hamburg. Clearly, the Rupelian deposits are not

only a product of a sea level highstand, but also a faultbound syntectonic feature. The depocenters are along or close to major fault systems striking N-S to NNE-SSW (which parallel the Upper Rhine Graben), however, the overall center of subsidence stretches from NW-SE (Fig. 4.4.21). In the eastern part of the basin, Oligocene sediments, which unconformably overlie the eroded MidPolish Swell, provide an upper time limit for the end of inversion-related uplift and erosion. Subsequently, since that time the subsidence pattern of the Polish Lowlands has not been related to the NW-SE-trending structures of the Trans-European Suture Zone, but rather to activity of the Carpathian forebulge and the depocenter, which developed to the N (Fig. 4.4.20d). Continued orogenic compression during the Early and Middle Tertiary modified the entire basin. Uplift led to significant erosion in Scandinavia and the Danish Basin with up to 1000 m amount (Japsen et al. 2002). The onset

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of erosion started already in the Neogene induced by tectonic movements prior to glacial erosion and post-glacial uplift. Major shifts from East to West in sediment transport directions are observed. Since then only transgressions of minor extent occurred during the Miocene and Pliocene, the marginal parts of the basin are characterised by paralic

liginite deposits and alluvial fans. Since the mid-Pliocene the North German Basin has reached again a continental stadium with limnic and fluvial deposits, however, these have been partially deformed by glacial tectonics and are overlain by deposits of the glacial period and post-glacial sediments.

4

Chapter 4.5

Glaciation, salt and the present landscape F. Sirocko · K. Reicherter · R. Lehné · Ch. Hübscher · J. Winsemann · W. Stackebrandt

4.5.1 Introduction The modern topography of the Central European Basin System (CEBS) is not only caused by the morphological shape of glacial landforms, e.g., moraines and ice-marginal valleys, and subsequent fluvial erosion, but clearly also reflects the pattern of the tectonic structures at depth, indicated by the preferred orientation of coastlines and rivers with three major tectonic strike directions: NW-SE, NNESSW and NE-SW (Fig. 4.5.1). Most likely, these pre-existing tectonic fault zones have fragmented the crust and lithosphere of the CEBS into distinct fault-bounded blocks. These block boundaries serve as discontinuities reactivated during times of isostatic adjustment in the major phases of the waxing and waning of Pleistocene ice sheets in the CEBS area. The effects of the isostatic rebound from the last (Weichselian) glaciation and subsequent crustal relaxation are still ongoing in Scandinavia and the southern Baltic and affect parts of the CEBS: Scandinavia is rising and most of the southern Baltic is actively subsiding. The loading/deloading processes must have been active also at the end of the earlier glaciations (Elsterian and Saalian) with maximum ice sheet extent over the CEBS.

ice sheets deep tunnel valleys were formed by subglacial meltwater erosion. The lithospheric forebulge formed in the periglacial regions, where permafrost reached depths of > 170 m. However, enhanced heat flow over the salt diapirs caused shallow permafrost depth compared to the surrounding sediments. Soils and unconsolidated Quaternary sediments of the subsurface were thus frozen into an irregular polygonal permafrost pattern with brittle (frozen) and plastic (unfrozen) patches with ice wedges, which must have provided a very inhomogeneous substrate for the advancing glaciers with implications for the abrasive power and stability of the ice sheet.

4.5.2 Modern topography and glacial isostasy

The frequent occurrence of pronounced lineaments in the modern topography indicates that unconsolidated Quaternary sediments only blanket the fault blocks. Expression of neotectonically active faults between blocks and along the major graben borders and salt diapirs are reconstructed from so-called fault-surface-penetrationpoints that are locally active still today and cause depressions (sink holes) with highest subsidence rates, partly leading to ground failures.

The modern topography of the CEBS landscape can be generalised into two major features in the digital elevation model (Fig. 4.5.1 a). The first is that of several arch-shaped moraine belts with a maximum elevation of more than 200 m stretching from western Poland to the North Sea and the Netherlands. These moraines are characterised by often laterally stacked thrust sheets and consist of unconsolidated till, glaciofluvial, glaciolacustrine and fluvial deposits. The base of the Quaternary deposits comprises mainly Tertiary clastic marine and coastal deposits. In eastern Germany extensive lignite deposits are intercalated in the sediments. In southern Lower Saxony push moraines partly overlie Cretaceous claystones, forming the basal detachment.

Several interconnected processes between salt and ice have shaped the glacial surface topography. The advancing ice sheets eroded the sediments and sedimentary rocks of the CEBS and interacted with salt diapirs at depth. Diapirs probably acted temporarily as barriers at which the glaciers stopped, depositing extensive meltwater sediments that subsequently became compressed by the advancing glacier, leading to the formation of large push moraines over the salt. Faults were locked under the ice cover, but the isostatic subsidence was compensated by a lithospheric forebulge in front of the ice sheet. Below the

The other dominating features are linearly orientated structures (Fig. 4.5.1 a). These orientations led Sirocko (1998) to express the hypothesis that the pattern of the North German rivers dominantly reflects the tectonic patterns of NW-SE and NNE-SSW striking faults at depth in the CEBS. These directions become visible in the shorelines of the Baltic Sea, but also reflect the course of river valleys like the Elbe, Weser and their tributaries. There are, however, also other directions such as the ice-marginal valley of Baruth, the rivers Havel and Spree. The nearly circular trend of river Havel, for example, is caused

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Figure 4.5.1. a) Digital elevation model of Northern Germany and Poland (data from Jarvis et al. 2006) showing major rivers, the maximum extent of the Weichselian, Saalian and Elsterian glaciations, after Ehlers et al. (2004). b) Salt diapirs and salt pillows of Northern Germany and Poland, after Lokhorst et al. (1998)

by a young subsiding block on top of the Elbe-lineament (Stackebrandt 2005). To explain the geodynamic cause of all these linear patterns it was necessary to quantify whether tectonic faults from the depth of the basin do indeed penetrate to the modern surface and/or if the position of the Quaternary inland glaciers was associated with tectonic units/structures at depth. This theory was proven by Reicherter et al. (2005), who showed by the evaluation of lineaments that basement faults in the supra-salt

Rotliegend (Fig. 4.5.2) are clearly reflected in the presentday landscape. The drainage pattern and the distribution of lakes in northern Germany correspond exactly to block boundaries in the deep basement and, hence, mark zones of subsidence and uplift. Additionally, the fluvial system in the CEBS is more complicated because of the occurrence of salt diapirs (Fig. 4.5.1b) and their response to ice loading. To take this complexity even further, a part of the system reacts diachronically. The Fennoscandian

Glaciation, salt and the present landscape

Ice Sheet in northern Germany during the last glaciation, the so-called Weichselian ice age some 20,000 years ago, reached a maximum ice thickness of 3000 m (Fig. 4.5.3), reconstructed from the mountain glaciation in Norway, isostatic rebound models and clay compaction of tills of the basal moraines (Piotrowski and Tulaczyk 1999). The load of the inland glaciers compressed the crust (or pressed the crust and lithosphere into the asthenospheric mantle), which after the retreat of the ice was unloaded and uplifted/expanded over several thousand years, attaining an isostatic equilibrium (Fig. 4.5.4). Accordingly, the prevailing regional tectonic stress caused by plate convergence in the Alps or spreading in the Central Atlantic is overprinted by glacial stresses (Roth and Fleckenstein 2001). The time of loading is considered to be long enough for a state of the isostatic equilibrium to be attained in the crust (Watts 2001). Elastic crustal flexure resulted in a circular depression below the centre, caused by radial outward flow of the asthenosphere (Daly 1934). The decay of the ice sheets must have changed the equilibrium situation in the crust and upper mantle. Important to note in this context is that the elastic response to fluctuating ice loads is regarded to be instantaneous, whereas the viscoelastic response of the mantle is much slower and must be active even thousands of years after the ice retreat (Stewart et al. 2000). Figure 4.5.2. Topographic map of the North German Basin; boundaries are shown in blue, major basement fault systems are shown in red; VDF marks the Variscan deformation front (from Reicherter et al. 2008)

Faults and diapirs must have acted during this transition as discontinuities concentrating displacements with offsets of tens to hundreds of metres as observed during the Holocene in Scandinavia (Fig. 4.5.5). There is at the moment not even a common hypothesis for the current geodynamic regime including regional tectonics, salt diaFigure 4.5.3. 3-D model of the Scandinavian ice sheet during the late Weichselian (LGM, last glacial maximum, approx. 20,000 years ago). The vertical scale is highly exaggerated. Data for model after Wu et al. (1999) and Siegert et al. (2001)

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pirism, glacial isostasy and modern landscape development for a reconstruction for the different glacial periods and interglacial periods of the past. This must remain a challenge for future dynamic modelling of the CEBS during the Quaternary. In the following chapter we will thus only present a few case studies to develop mechanistic model explanations.

ity, which led to major earthquakes (Magnitude > 8) and surface ruptures in northern Sweden and Finland in the early phase of the uplift (Mörner 2003). The deloading accompanying geological effects in the peripheral areas during the Holocene were not as dramatic as in the centre of the ice load. However, differential subsidence and/or uplift must have been on the order of several mm per year and in the beginning with much higher rates.

4.5.3 Crustal movements, seismicity and landscape formation

The complex seismic strain-release patterns modified during the decay of the ice sheets have been termed “deglaciation seismotectonics” by Muir-Wood (2000). Measurable crustal deformation is still the consequence of the mantle response to deglaciation (Scherneck et al. 1998), accompanied by decelerating seismic activity (Mörner 2003). It has to be pointed out that it is generally very difficult to distinguish between ice-induced earthquakes and earthquakes resulting from plate tectonics in areas of repeated glaciation/deglaciation cycles.

The adjustment of the crust and lithosphere over the CEBS into an ice-free equilibrium state must have been accompanied by several geological phenomena, like seismicity and earthquakes (Fig. 4.5.5 a, after Gregersen and Voss 2008), with the formation of large faults and surface ruptures, slope failures and liquefaction as well as raising shorelines (Stewart et al. 2000; Mörner 2003). The isostatic uplift due to the postglacial rebound in Fennoscandia is also associated with moderate to major seismic-

The shorelines of the early Holocene Baltic Sea have been uplifted by at least 300 m since the deglaciation.

Figure 4.5.5. a) Recent uplift and subsidence in the North German Basin area in mm/year (modified after Frischbutter et al. 2001) b) Postglacial seismicity and active faults in Scandinavia (from Gregersen and Voss 2008)

Glaciation, salt and the present landscape

Box 4.5.1 Glaciations The Central European Basin was affected by 3 major glaciations during the Elsterian, Saalian and Weichselian periods. Figure 1 shows the very latest stratigraphic information, compiled by Litt et al. (2007). Sediments of the oldest Quaternary glaciation, the Cromerian Complex, are rarely known. Studies on the number and extent of ice sheets transgressing the Central European Basin are mainly based on the interpretation of glacial landforms and/or the distribution of glacial deposits (Ehlers et al. 2004). Tills are separated by glaciolacustrine and glaciofluvial deposits. In front of the ice sheets numerous large ice-marginal lakes formed capturing glacial meltwater and surficial water due to the blockage of natural drainage pathways to the north by ice sheets. Widespread fluvial erosion took place during interglacials and coastal areas became flooded due to rising sea-levels (e.g., Caspers et al. 1995; Eissmann 2002; Ehlers et al. 2004; Winsemann et al. 2007a,b).

Modern geodetic measurements still reveal uplift rates of up to 11 mm/year (Fig. 4.5.5a; James and Lambert 1993; Scherneck et al. 1998; Milne 2001). Lower Saxony and most of Schleswig-Holstein of Northern Germany are instead actively subsiding at rates of up to 1.5 mm/year (Fig. 4.5.5b). The reason for subsidence is asthenospheric back flow towards the centre of the former glaciation in Scandinavia and isostatic crustal re-adjustment in the CEBS area. Apparently, the regional tectonic stresses from the Alpine front and the Mid-Atlantic ridge push gain again more influence as pointed out by stress distribution modelling in the previous chapter.

Major basement faults in Northern Germany are oriented NW-SE, while minor faults trend NE-SW and NNE-SSW (Reicherter et al. 2005). The first indication that these orientations of structures in the deep crust are also expressed at the surface (Fig. 4.5.1) came from Hennig (1906). A similar interpretation of the orientation of rivers related to tectonic lineaments has been pointed out by Sirocko (1998) in particular for the river courses of the Oder, the Elbe (only the lower course from Magdeburg to the estuary near Hamburg) and the Weser, which either follow NW-SW or NNESSW directions. The location and trend of lineaments and faults correspond to old structures in the Variscan and pre-

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Variscan basement (Fig. 4.5.2). The digital elevation model exhibits this observation, which can also be applied for other parts of the landscape forming (Fig. 4.5.1). In particular, the shorelines of the Baltic Sea are relatively linear, but mostly with SW-NE orientation which is the third major tectonic orientation of the CEBS. The river Oder between Germany and Poland marks a sharp change in the preferred orientation. The Polish rivers and shores are dominated by the SW-NE orientation, whereas in Germany, NW-SE or NNE-SSW orientation is dominating. These three directions dominate the surface topography of the entire CEBS. A strong indication that the modern landscape is at least partly a mirror of the basin history is the continuation of the Polish salt pillows at depth and the modern Baltic Sea coastline southwest of Rügen Island (Fig. 4.5.1b). However, it is unlikely that such surface lineaments are directly forced by processes from below (endogenic or halokinetic), but much more likely these lineaments represent block boundaries, activated by the repeated loading and unloading under the glacial ice masses. Modern rivers and shorelines still mark the boundaries between individual blocks.

Inland ice masses during previous glaciations had an even larger extent than during the Weichselian, when glaciers were restricted to the area north of the river Elbe. The Saalian and Elsterian glaciation (ice margins in Fig. 4.5.1, see Box “Glaciation” for chronology) reached further south up to the foot of the Harz Mountains and into the Netherlands. Maximum ice sheet thickness and, hence, ice load during these two earlier glaciations were higher over the CEBS area than during the Weichselian, leading to strong ice-isostatic effects. The model of the ice sheet – crust interaction (after Mörner 2003) indicates that under the conditions of the Elsterian or Saalian there was an enhanced ice loading of the CEBS sediments. This effect was accompanied by the development of a crustal forebulge extending up to several hundred kilometres distance from the ice front in the distal foreland (Fig. 4.5.4). Hence, the effects of isostatic rebound were not only limited to the former ice-covered areas, but also to the periglacial landscape in front of an ice sheet.

Figure 4.5.6. a) Modern topography of the landscape around Rederstall (Schleswig Holstein) in parallel orientation to the geological strata below (Lehné and Sirocko 2007), data from Baldschuhn et al. (1996). b) to d) Schematic sketch of a forward moving glacier with a salt diapir at subsurface

Glaciation, salt and the present landscape

Figure 4.5.7. Ice load induced tectonism (modified from Liszkowski 1993) 1a-c Influence of ice-loading/unloading on a normal fault with crustal failure and reactivation of inherited structures. Red arrows indicate relative crustal movements 2a-c Effects of ice-loading/unloading on a pre-existing graben system with conjugate faults 3a-c Consequences of ice-loading/unloading on faults and salt dynamics associated with a salt diapir, red arrows indicate relative diapiric movements (Liszkowski 1993)

4.5.3.1 Regional and case studies Schirrmeister (1998) demonstrated that most Weichselian moraines in Mecklenburg-Vorpommern and Brandenburg are in direct contact with salt structures at depth. Lehné and Sirocko (2007) reported this relationship also for a section through the salt structures of Hennstedt and Tellingstedt south of Husum in Schleswig-Holstein (Fig. 4.5.6 a) where a parallel orientation of the salt structures with the geological strata above and the modern topography becomes evident. The Quaternary deposits along the section are of Saalian age, which would invoke an average continuous uplift of 0.3 mm/year if the surface structures would be caused by rising salt alone. However,

the pattern of parallel topography and salt at depth could be explained also by another scenario, which starts with the glacier advancing towards a salt structure leading to an additional load on the subsurface, creating local accommodation space for meltwater deposits in front of the rising salt diapir. The meltwater deposits subsequently became compressed, leading to the formation of a push moraine. Such a model would help to explain why glacial push moraine belts are often (but not always) located immediately north of a salt diapir. Other mechanistic models for salt – ice interaction have also been proposed (Fig. 4.5.7). An attendant and very important circumstance of ice loading is stabilisation of

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Figure 4.5.8. a-c) Seismic profile of a salt diapir with near surface faults and fault projection to the surface. d) Map of Schleswig Holstein with fault penetration points, determined by evaluation of 454 seismic lines of the oil industry (from Lehné and Sirocko 2007)

Figure 4.5.9. 3-D Plot of the base of Late Cretaceous in the lake Plön area, additional lines for faults of the eastern boundary of the Glückstadt Graben, position of an Elsterian tunnel valley and position of project area Klein Neudorf

Glaciation, salt and the present landscape

Box 4.5.2 Permafrost The presence of permafrost in central Europe during various glacial stages of the Pleistocene is indicated by relict ice-wedge structures, which have been frequently observed in excavations of lignite mining districts of East Germany (Eissmann 2002) and are observable even today in vegetation patterns, mimicking relict ice-wedge polygons in northern Germany (Delisle et al. 2007). However, no direct indication of the former permafrost depth is known from the sub-surface sedimentary record. Figure 1 shows the calculated variations in depth of permafrost along a 75 km long profile crossing several salt structures of the Glückstadt Graben. The permafrost reconstruction is based on the mean annual ground temperatures curve (MAGT) for north-central Europe, the marine proxy record of ODP-site 659, the average porosity of the sediments and the average thermal conductivity. The model predicts a maximum permafrost depth of about 170 m during the last million years. Above salt structures the depth of permafrost is reduced by up to 40 m. Higher thermal conductivity of rock salt results in enhanced heat flow through salt structures at the expense of the surrounding country rock. Therefore, enhanced heat flow over salt domes impeded the development of permafrost, and reduced heat flow in the adjacent country rock favoured deeper penetration of the lower permafrost boundary. The interplay of salt diapirism and fault movements in the North German Basin is difficult to assess for several reasons: during permafrost times the groundwater was frozen and not able to dissolve or modify evaporites (e.g., anhydrite to gypsum); due to the ice load faults and rising diapirs were blocked: the effective stresses were the sum of the regional tectonic stresses and the ice-induced stresses.

faults under the ice, leading to the suppression of seismicity and a decrease in fault stability beyond the ice margin (Johnston 1987; Johnston et al. 1998). Ekström et al. (2003) detected dozens of previously unknown, moderate earthquakes beneath large glaciers. These are relatively “slow”, because the driving mechanism is thought to be wave-like glacial movements they are termed “glacial earthquakes”. However, loading by small ice sheets (radius approx. 300 km) causes an increase in stability of less than 1 MPa at shallow depth, but promotes instability at greater depths (Stewart et al. 2000). In contrast, unloading also decreases stability relative to the initial pre-glacial state (Fig. 4.5.7 1a-c). Repeated ice progression during

interstadial and interglacial periods must have significant influence on the fault activity, i.e., faults get locked under the ice loads and are reactivated/initiated in front of the ice mass (Fig. 4.5.7 2a-c). The influence of underlying salt (which is regarded as incompressible) on fault activity in the vicinity of diapirs is shown in figure 4.5.7 3a-c. During glaciation the rise of the diapir is hindered, and faults are blocked. After the retreat of the ice sheets, faults are reactivated. To verify or falsify these model assumptions we examined whether faults are capable to reach the surface even through unconsolidated Quaternary deposits of some-

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Figure 4.5.10. Base of Quaternary deposits (deepest sections depicting the tunnel valleys), reproduced from Stackebrandt et al. (2001)

Glaciation, salt and the present landscape

Figure 4.5.11. a) Map of Ostholstein with salt diapirs and salt pillows and the two project areas of Lake Plön and the offshore Gabelsflach, both along the very eastern flank of the Glückstadt Graben. b) and c) Photo of an actively sinking depression near Klein Neudorf, east of Lake Plön. Aerial picture of the enlarging depression in 1959 and 1999 (from Lehné and Sirocko 2007)

times several hundred meters thickness. The approach is based on the observation that an observed offset at depth multiplied by a factor of 20 approximates the minimal continuation of a fault (J. Urai, pers. comm.). The fault at 500 m depth in the example of figure 4.5.8 has an offset of 40 m and will continue for at least another 800 m. Accordingly it has the potential to reach the surface and cause a penetration point. This approach was used for a systematic mapping of all available seismic lines in SchleswigHolstein and resulted in a map of penetration points (Fig. 4.5.8, from Lehné and Sirocko 2007), which shows that most penetrations points are associated with salt diapirs that have risen to depths of a few hundred meters below the surface. There are, however, also clusters of points in areas not associated with salt diapirism, for example north of the most eastward salt wall. This salt wall north of Bad Segeberg parallels the eastern flank of the Glückstadt Graben and is associated with deep reaching faults down to the base of the CEBS, in particular east of Lake Plön. A digital elevation model for the Lake Plön area reveals the base layer of the Late Cretaceous (data after Baldschuhn et al. 1996; Fig. 4.5.9). The eastern shoreline of the lake is above two parallel running faults striking in a SSWNNE direction, associated with the deep Segeberg-Plön salt wall / pillow. Both faults are normal faults with a displacement of almost 2000 m with respect to the Glück-

stadt Graben. The orientation is remarkably parallel to the eastern shore of Lake Plön. The western shoreline of the lake is also straight and parallels a small Miocene graben. Accordingly, Lake Plön can be regarded as a tectonic lake and not only a basin formed by glacial erosion. However, the shape of the modern Lake Plön matches also an Elsterian tunnel valley. These valleys of up to 500 m depth developed at the base of the Elsterian ice sheet. Figure 4.5.10 reproduces a compilation map of Stackebrandt et al. (2001) and Stackebrandt (2004). The orientation of the tunnel valleys is on average NE-SW and thus not fully identical to the NNW-SSE strike. The tunnel valleys were mainly formed by subglacial meltwater erosion (Piotrowski 1997; Huuse and Lykke-Andersen 2000; Jørgensen and Sandersen 2004) and the direction of tunnel valleys in general is thought to be perpendicular to the ice sheet margins, documenting the flow direction of the ice. It is discussed since long whether the subglacial drainage pattern could have partly followed old tectonic structures in the subsurface. The modern Lake Plön is thus most probably an inherited structure, which originated in the Miocene under intense tectonic stresses due to the Alpine orogeny and maximum uplift rates of salt diapirs in the CEBS, it would thus be an example for the model explanation of figure 4.5.7 case 2.

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Figure 4.5.12. a) Seismic profile along the Gabelsflach in the southwestern Baltic Sea (Fig. 4.5.1a) located above a salt structure. b) Line drawing of parametric echosounder data (/Innomar SES system/) across the “Gabelsflach Heigh” with Holocene sediments superimposing an unconformity from glacial erosion. c) Vertical succession of pre-tectonic (parallel) and syn-tectonic (divergent) Holocene sediments at the flanks

The larger Lake Plön area also provides evidence that these faults along the eastern flank of the Glückstadt Graben are still active today (Lehné and Sirocko 2005, 2007). This inference comes from a recently enlarging depression near Bosau/Klein Neudorf. The depression is located directly above the large faults, which offset the Late Cretaceous by some 2000 m (see location of depression in Fig. 4.5.9). The depression had a diameter of 15 metres in 1959 (Fig. 4.5.11, yellow line); in 1999, 40 years later, the diameter had enlarged to 65 metres (Fig. 4.5.11, blue line). The amount of subsidence was more than one metre in 50 years, thus on the order of several cm/year, which is much higher than the general basin subsidence of less than 1 mm/year. The high sinking rates require a geodynamic process operating on the order of cm/years. Such rates have been reported from local maxima of dissolution

of salt in the subsurface or from compacting peat. A 65 m borehole into the depression, however, has neither revealed peat nor salt at depth, but only sand. This corroborates the very young nature of the depression, because if this had been a late Weichselian kettle hole, the depression would have been filled with Holocene gyttja and peat. Dissolution of salt in the subsurface is also impossible because the salt surface is at 2000 m depth where pore water is saturated with salt. The depression, however, lies exactly above steep faults with high slip rates at depth. If these faults are associated with extensional stress and a fault penetrates through unconsolidated sediments, a local sinking might well be the expression of a surface penetration point. This is the most likely interpretation for the depression at Klein Neudorf and for many other local depressions in the surface sediments of the tectonically active CEBS, today and

Glaciation, salt and the present landscape

Box 4.5.3 Salt diapirs at the surface; the Sperenberg example Salt diapirs have penetrated the modern surface at only three locations in the CEBS: at Bad Segeberg in Schleswig Holstein, Lüneburg in Lower Saxony, and Sperenberg, south of Berlin. Figure 1 shows an elevation model for Sperenberg (Stackebrandt 2005), which is situated on a NW-SE striking neotectonically active fault zone. The caprock consists of several tens of metres thick gypsum and anhydrite; the salt has been encountered 120 m below the surface. The Sperenberg diapir was only loaded by ice during Elsterian and Saalian glaciations. Thus, it has been ice-free for about 130.000 years. The Saalian tills, which cover the structure, were uplifted differentially. During the subsequent Weichselian ice age, Sperenberg formed a topographic high, possibly related to a Nunatak. The isostatic caused uplift of the diapir leads to salt solution in the sweet water reaching neck of the diapir characterised today by circular lakes, indicating active present-day landscape sculpturing processes.

also during the past, when isostatic movements must have affected fault activity in northern Germany, in particularly during the early deglaciation. The paragraphs above describe examples of tectonic and halokinetic activity on land. The following section will now evaluate if these mechanisms also operate offshore. We studied in detail the Gabelsflach, which is a bathymetric high of less than 10 m water depth in the southwestern Bay of Kiel, located on the continuation of the Glückstadt Graben border from the area of Lake Plön to the north into the Baltic Sea (Fig. 4.5.11 a). Hübscher et al. (2004) describes marine seismic surveys, which reveal a deep salt diapir with pronounced bathymetric changes on the sea floor above (Fig. 4.5.12). The outcropping sediment on top of the salt diapir is a glacial till, whereas at the flanks we find laminated mud. The western flank of the uplifted sea bottom is subdivided into a parallel lower and a divergent upper succession. Parallel layers were deposited during tectonic quiescence, whereas divergent reflection points towards vertical movement, which confirms that

the entire Gabelsflach structure was uplifted during the later Holocene. Thus, the salt diapir below the Gabelsflach structure was active during postglacial times. This fact proves that also halokinetic forces were active during the Holocene, not only isostatic processes. The picture of an active landscape formation from endogenic-exogenic forces in the CEBS is based on GIS analysis of the topography, models of isostasy and rebound, reconstruction of fault activity and modern subsidence rates. A common interpretation of all observations highlights the role of the Quaternary glaciations, which flexured the crust and fractured it into blocks along existing faults at depth that reach deep into the basin. Salt diapirs were halokinetically rising mainly during the Tertiary, but are still active during the Holocene. Today, even after the decay of the ice sheets, ongoing subsidence still reflects the long-lasting isostatic adjustment of the crust and lithosphere in the CEBS, a process which had shaped the modern landscape in particular during the final phase of the last glaciation.

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Chapter 5

Salt dynamics

5

Chapter 5.1

Salt as sediment in the Central European Basin system as seen from a deep time perspective J.K.Warren

5.1.1 Introduction This section places salt (halite-dominant) sediment found in the Central European Basin system into a world-scale understanding of similar evaporitic sediments. The thickest and most laterally extensive salt beds (100-300m depositional thickness) in the Central European Basin system were deposited and stacked into the intracontinental Permian Zechstein salt group that is up to 2-3 km thick, while thinner less extensive beds halite (10-50 m thick) characterise parts of the Triassic succession with salt beds stacking into evaporite sequences up to 200-300 metres thick (Fig 5.1.1). Depositional setting at the time of Zechstein salts units precipitated was a hydrographically isolated marine-fed intracontinental sag basin locally subject to rifting and basement extension (Figs. 5.1.1a, 5.1.2a). The depositional setting for much of the Triassic was a continental interior basin with relatively minor evaporite deposition over much of the basin interior for much of the time. On three occasions in the Triassic the basin became a site of bedded halite deposition (Röt (Buntsandstein),

Muschelkalk and Keuper halites) with the depositional area becoming smaller with each succeeding cycle of subsidence (Figs. 5.1.1b, 5.1.2a; Michelsen and Clausen 2002). The degree of depositional “marineness” during halite times increased eastwards towards Poland and local depocentres were related to subsidence of the Central and Horn grabens and the Glückstadt and Polish troughs (Glennie 1986). Triassic halite was also deposited in other smaller basins to the west and south of England, as well as over a post-Variscan zone of weakness on continental Europe, known as the Hessian Trough. The various Triassic halite beds tended to be thicker in actively subsiding rifts about the marine-influenced edges of this huge continental interior desert basin (Figs. 5.1.1b, 5.1.2a). Triassic extension that drove the subsidence also drove widespread diapirism and namakier formation sourced in the underlying Zechstein salt. Evaporitic sediments of Zechstein Group are subdivided into Z1-Z6 units based on clay-carbonate anhydrite-halite groupings (Fig. 5.1.2b). This classic subdivision has

Figure 5.1.1. A) Equal area palaeogeographic reconstruction for the Late Permian Zechstein Basin (250 Ma) showing palaeolatitude position of 30N (after Torsvik et al. 2002). B) Regional Triassic palaeogeography of NW Europe (in part, after Ziegler 1990)

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been somewhat modified by more recent sequence stratigraphic approaches to the stratigraphy (Tucker 1991; Strohmenger et al. 1996). Halite dominates the basin-centre positions and units Z2 and Z3, which are also the halite successions most prone to halokinesis (Fig. 5.1.2c). Classically, the Triassic is subdivided into three based on lithology (von Alberti 1834): (i) the predominantly continental Buntsandstein, (ii) the marine Muschelkalk, and (iii) the mixed marine to continental Keuper (Fig. 5.1.2a).

Worldwide, bedded salt sediments or their residues are a depositional response to a spectrum of chemical and biochemical arid-zone settings ranging in age from the Archean to the present (Warren 2006). Ancient salt beds were deposited under climates and hydrologies that are typically interpreted using the concept of uniformitarianism (Hallam 1990). In a uniformitarian conceptualisation it is important to draw the distinction between strict actualistic uniformitarianism and a more prosaic or realistic uniformitarian approach that takes Figure 5.1.2. Halite Stratigraphy. A) Chart of halite-entraining stratigraphic units and regional tectonic events (after Mohr et al. 2005). Note that times of salt deposition correspond to times of regional extension. B) Zechstein stratigraphy in Germany (after Strohmenger et al. 1996). C) Classic stratigraphy - lithology distribution from basin edge to centre (after Geluk 2000)

Salt as sediment in the Central European Basin system as seen from a deep time perspective

into account longer-term fluctuations in earth scale processes. Consider current views of the state of the world‘s climate as an example of this distinction in the application and significance of uniformitarianism. The popular anthropocentric view of today‘s climate, especially in the popular press, tends to be a strict actualistic view, where world climate, without interference from mankind, is viewed as a constant, with its invariability of great benefit to humankind. From a geologic viewpoint, which takes a longer time view for climatic comparison, we know that climate changes at times and rates that are outside the very limited ranges observable in terms of a human life span and even outside the range of variation sampled by the time frame of human evolution (Hay et al. 1997). This geological perspective is sometimes called the view from “deep time.” Using sediment distributions through time as their ground truth, the climate modelling of Sellwood and Valdes (2006) shows that dense Mesozoic forests grew close to both poles but experienced months-long daylight in warm summers and months-long darkness in cold snowy winters. Ocean depths were warmer (8 °C or more to the ocean floor) and reefs, with corals, grew 10° of latitude further north and south than at the present time. The whole Earth was warmer than now by 6 °C or more, giving more atmospheric humidity and a greatly enhanced hydrological cycle. Much of the rainfall was predominantly convective in character, often focused over the oceans and leaving major desert expanses on the continental areas. Permanent polar ice sheets are unlikely to have been present because of the high summer temperatures achieved and, within appropriate latitudinal ranges (the horse latitudes), evaporite deposition occurred in this greenhouse time at a scale and diversity not seen in today‘s icehouse climate mode. Evaporites in the deep time framework of earth history are a time-variant system related to complex feedbacks between atmospheric circulation and chemistry, seawater composition, seafloor spreading rates, subduction and continent-continent interactions (Warren 2006). Accordingly, the interpretive approach in this chapter is based on uniformitarianism using a deep time perspective at a world-scale resolution. The salt focus throughout this chapter is halite (not sulphate). We shall consider how it can be deposited as thick laterally extensive beds (a.k.a. saline giants or megahalites) and why such deposits formed best at particular times in the past. We shall not consider modern or ancient sulphate systems in any detail; for an interested reader, Becker and Bechstädt (2006) is a good introduction to the

Zechstein sulphate platform literature. By chapter end, we shall come to understand why no marine-fed megahalites are currently accumulating anywhere in the world today. We shall first look to the Quaternary to establish boundary conditions for the appropriate set of processes for salt sediments in each of the topic sets of this paper; climate, brine evolution and basin-scale depositional systems. But we shall look at each set of Quaternary-defined processes from a deep time perspective and show how boundary conditions may change. Such an approach underlines limitations of strict actualism when interpreting many ancient salt beds. Finally, we shall apply this approach to improve our understanding of namakier and depositional setting in Triassic salt beds of the Central European Basin system.

5.1.2 Mother brines: isochemical systems? To deposit a thick evaporite, mother brine must concentrate via solar evaporation in a hydrographically-isolated depression on the earth‘s surface where there is the potential for more water to leave the basin via evaporation than enter. Thick salt beds require the brine body to remain within a hydrological setting that is stable over a suitable time frame of tens of thousands to hundreds of thousands of years (Warren 2006). There must also be a sufficient volume of mother brine entering the basin to allow a thick salt mass to accumulate and be preserved. Mother brines can be marine, nonmarine or hybrids (mixed marine and nonmarine including resurgent hydrothermal and basinal waters). In this chapter we concentrate on chemistry of marine-derived brines as their ancient counterparts were the dominant brine source for all thick and widespread halite accumulations, including the Zechstein and Triassic salt beds of Central Europe. Same-scale comparison of depositional settings worldwide show meteoric waters are never volumetrically sufficient to act as ionic reservoirs for largest ancient saline giants, although they are the dominant source brine for numerous moderate-sized lacustrine salt successions (Warren 2006). To interpret brine-related controls on evaporite sediments we first use modern seawater as a guide to the sequence of minerals that precipitate as marine brine concentrates. Throughout the Quaternary, composition and ionic proportions in a seawater-derived brine have been near constant and its ionic composition dominated by Na-Cl with lesser Mg, SO4, Ca and HCO3 (Fig. 5.1.3a). Likewise, the salt suite formed by simple single-stage evaporation of Quaternary seawater is predictable and well understood. Almost all of the bicarbonate in concentrating seawater is used during the initial precipitation of

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J.K. Warren Figure 5.1.3. A) Changes in ionic proportions due to sequential precipitation of aragonite, gypsum, halite and bittern salts as modern seawater is evaporated to 100 times its original concentration (in part after McCaffrey et al. 1987). B) Pre-bittern (0-400‰) mineral precipitation sequence in a concentrating marine brine (replotted from Table 1 in Briggs 1958)

mesohaline carbonates, when salinities are some 1.5 to 3 times that of the original seawater (>60‰). Extraction of bicarbonate leaves excess Ca in the concentrating brine, which then combines with sulphate to precipitate gypsum or anhydrite. This begins when the seawater brine reaches 5-6 times its original concentration (150‰; Fig. 5.1.3b). Ongoing gypsum/anhydrite precipitation removes almost all of the Ca from the brine leaving excess sulphate in the liquor. After halite has precipitated (onset at 10-11 times

original seawater concentration; ≈350‰), the sulphate combines with the magnesium (conservative concentration to this point) to precipitate a series of MgSO4 salts along with the KCl bitterns. The bitterns begin to precipitate when seawater reaches 60-70 times its original concentration. The volume of potential halite precipitate held as solute in seawater far exceeds that of any other salt. When

Salt as sediment in the Central European Basin system as seen from a deep time perspective Figure 5.1.4. Changes in Phanerozoic seawater chemistry. A) concentration of Mg in fluid inclusions in primary chevron halite. A value of 35 mg/l is chosen -pink line) to separate MgSO4 oceans from CaCl2 oceans (after Zimmermann 2000; Horita et al. 2002). B) Variation in sulphate levels in Permian and Triassic chevron halite from selected European deposits and elsewhere (after Kovalevych et al. 2002b). C) Salinity change over the last 500 Ma (after Hay et al. 2006)

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seawater brine is evaporated completely, halite by volume makes up more than 90% of the total amount of precipitated salts (Fig. 5.1.3b). However, any evaporite basin can attain hydrologic equilibrium (amount of water lost via evaporation equals the amount of solute inflow) at any stage of the evaporation series. If hydrologic equilibrium is attained during salinities appropriate for gypsum precipitation (≈135-270‰) then the basin fill will be dominated by anhydrite, not halite. Complete evaporation of Quaternary seawater produces two sequential types of high density, high viscosity brine: Na-Cl brines and Mg-Cl-SO4 brines. Na-Cl brines dominate in the salinity range 35-330‰. Brines at higher salinities become progressively dominated by Mg-Cl-SO4, indicating preferential removal of Na by halite precipitation. At first, Mg is the dominant cation in the bittern brines but, as the various MgSO4 bitterns crystallise, K takes its place and the various KCl bitterns precipitate. Cl remains the dominant anion throughout the evaporation sequence of modern seawater. Oceanic chemistry shows variation in both its ionic proportions and salinity over a longer Phanerozoic time frame (Fig. 5.1.4; Horita et al. 2002; Zimmermann 2000; Kovalevych et al. 1998; Hay et al. 2006). Unlike the carnallite- and MgSO4-dominated bittern salts precipitated from modern seawater, the bittern assemblies in many Phanerozoic marine-fed salt basins are dominated by sylvite and the sediment lacks primary MgSO4 salts. MgSO4depleted evaporites indicate times in the Phanerozoic when the world oceans were Ca-rich and SO4-poor (CaCl2 seas), compared to Quaternary seawater (Fig. 5.1.4a). These CaCl2 oceans correspond to what are also known as the Phanerozoic calcite oceans where calcite and Mg-calcite, not aragonite, were the dominant marine carbonate minerals. CaCl2 seas coincide with periods of rapid seafloor spreading, times of high influxes of mid-ocean-ridge brines rich in CaCl2, and elevated sea levels (Hardie 1996). Thus Phanerozoic oceans have undergone two long-term fluctuations in major-ion composition (Mg/Ca ratios) that are in phase with 100-200 million-year oscillations in sea level, icehouse-greenhouse climates, and global volcanicity. Within this deep time framework, fluid inclusion studies of Permian halite chevrons have been done by Kovalevych et al. (2002a) using samples from several Permian evaporite basins of East and Central Europe (Asselian and Sakmarian Dnipro-Donets Basin of Ukraine, Sakmarian Dvina-Sukhona Basin of Russia, Kungurian Solikamsk Basin of Russia, Zechstein Basin of Poland and Peri-Baltic Russia). A marine origin of the Permian salt in most of their studied sections was supported by isotopic composition of sulphate sulphur. Their work showed synpre-

cipitational brines were of Na-K-Mg-Cl-SO4 (SO4-rich) chemical type, revealing chemistry similar to Quaternary seawater, although sulphate levels were somewhat lower than in today‘s ocean (Fig. 5.1.4b). Such general similarity in ionic proportions is to be expected as, like today, the Permian was another time of aragonite seas (Hardie 1996). A similar inclusion study of the Triassic Röt Halite by Kovalevych et al. (2002b) shows that the early Triassic Röt brines are also Na-K-Mg-Cl-SO4 (SO4-rich) waters, i.e., another set of ionic proportions similar to that of modern seawater (Fig. 5.1.4b). However, the Röt brines differ from the composition characteristic of modern concentrated seawater by a decrease in the content of sulphate ion and an increase in the content of potassium ion. Also, the relative content of sulphate ion in the Early Triassic Röt Basin was slightly lower than that in Permian evaporite basins of Europe, but higher than that in Lower and Middle Permian and Middle and Late Triassic evaporite basins (Fig. 5.1.4b). The implication is that the content of sulphate ion in seawater was subject to changes in Permian and Triassic times and that the decrease in sulphate from the Permian to the Triassic indicates the onset of the transition from a MgSO4 ocean into the CaCl2 ocean that characterised much of the period from the Triassic until the Miocene. Inclusion studies of chevron halite also show that independent of Phanerozoic fluctuations in Mg, Ca and SO4 proportions, the ionic composition of seawater has always been dominated by Na and Cl. Halite and/or gypsum (now anhydrite in the subsurface) have always been the most volumetrically significant marine evaporite salts. Yet over the same Phanerozoic time frame the salinity of the oceans shows a general decrease (Fig. 5.1.4c; Hay et al. 2006). The greatest changes in salinity are related to major extractions of salt into the young ocean basins. The last major extractions of salt from the ocean occurred during the Late Miocene, shortly after the large-scale extraction of water from the ocean to form the ice cap of Antarctica. Accordingly, salinities during the Early Miocene were between 37‰ and 39‰. The Mesozoic was a time of generally declining salinity associated with the salt extractions in the opening North Atlantic and Gulf of Mexico (Middle to Late Jurassic) and South Atlantic (Early Cretaceous). As we shall see in a later section, the earliest of the major salt extractions of the Phanerozoic occurred during the Carboniferous-Permian and related to the final accretion of the Pangaea supercontinent. Other than in the earliest Cambrian, there were few episodes of salt extraction prior to the Carboniferous. The Hay et al. (2006) model suggests that much of the Palaeozoic was a time of relatively stable but slowly increasing salinities (in parts per thousand) ranging from the lower to the upper 40’s.

Salt as sediment in the Central European Basin system as seen from a deep time perspective

This change in salinity implies that arid-zone marine-fed evaporites formed more readily from seawater brines in the early Phanerozoic epeiric seas compared to the less saline oceans and narrower steeper shelves of today. It also implies that halite saturation was more easily attained from Palaeozoic ocean waters, yet for much of this time the volumes of extracted halite were low (Fig. 5.1.7). The Hay et al. (2006) salinity curve shows the oceans were at their freshest in the Late Cretaceous some 80 Ma, not today. This is because a substantial part of the Mesozoic salt mass, deposited in the megahalites of the circum-Atlantic and circum-Tethyan basins, has since been recycled back into today‘s ocean via a combination of dissolution and halokinesis. The significance of this salinity change in relation to Phanerozoic plate tectonics and the volume of evaporite sediments are discussed further once we have developed an understanding of the significance of climate to the deposition of salt sediments.

5.1.3 Evaporite sediments and climate Modern continental salts typically accumulate within ground-water discharge regions in semi-arid to hyperarid

deserts (Fig. 5.1.5a). Coastal evaporites occur at the oceanic edge of the same deserts, typically in coastal depressions fed by marine seeps or by rising groundwaters along coast-parallel mudflats. Evaporites also form lake precipitates and efflorescences in cold polar deserts in Antarctica, but the volumes of Holocene salts in these cryogenic regions pale to insignificance compared to modern arid settings closer to the equator. Climate drives deposition of evaporite sediments at scales ranging from broad latitudinal belts (100s km) down to microclimates (cms). At the broadest scale, evaporite sediments form in deserts and are the result of largescale atmospheric circulation. Modern deserts cover more than 30% of the world‘s landsurface, mostly within two belts lying 15° to 45° north and south of the equator (Fig. 5.1.5a and inset in Fig. 5.1.6). Belts of aridity sit beneath cold dry descending air masses of the Hadley Cells, which define high-pressure belts known as the subtropical horse latitudes. Global-scale atmospheric circulation is driven by varying intensities of solar irradiation, which is most intense directly above the equator and lessens toward the poles (Fig. 5.1.5b). Hence, the equatorial belt experiences greater insolation than the adjacent temperate latitudes. Equatorial air warms as it rises, creating a tropical belt of

Figure 5.1.5. Factors involved in creation of the world’s deserts. A) World distribution of modern deserts as determined by plotting areas with less than 250 mm annual precipitation. 1. Australian, 2. Great Basin, 3. Mojave (Sonoran), 4. Chihuahuan, 5. Baja California, 6. Peruvian, 7. Atacama (Chilean), 8. Patagonian, 9. Sahara, 10. Namibian, 11. Kalahari, 12. Arabian, 13. Turkestan, 14. Iranian, 15. Thar, 16. Gobi. Also shows polar regions with less than 25 cm precipitation. B) Longitudinal cross section through the earth’s atmosphere showing major circulation cells. Belts of cool dry descending air at 30 °N and S of the equator create the main arid zones of the world (shaded). C) Adiabatic or rain shadow desert where rain tends to fall on the seaward side of the mountains. D) Coastal deserts form in response to cool upwelling ocean waters and the cool dry onshore winds warming as they pass overland (after Warren 2006)

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Table 5.1.1 World-scale climate processes controlling distribution of evaporites in modern deserts (after Warren, 2006) Cause and scale

Situation

Quaternary examples

Descending or stagnant air mass (Planet scale)

Subtropical regions. Regions beneath the descending limbs of Hadley cells (centred around 20-30° north and south of the Equator.

Sahara, Namibia, Kalahari (Africa) Chihuahuan and Sonoran (North America) Great Sandy and Simpson deserts (Australia) Arabian, Iranian, Thar deserts (Middle East and India)

Continentality (Plate tectonic scale)

Regions within the continental interior. Locations isolated by distance from moisture sources, principally the oceans.

Saharan (Africa) Turkestan, Talimaklan, Gobi (Asia) Great Basin (North America)

Adiabatic or Rainshadow (Plate tectonic scale)

Regions located in the lee of topographic barriers, such as mountains and plateaus. Cooling of rising airmass creates rainfall on the windward/ seaward side of a mountain range or barrier. As airmass subsides on the leeward side it warms and dries.

Turkestan, Talimaklan (Asia) Great Basin, Colorado Plateau (North America) Monte, Patagonian (South America)

Cool sea surface temperatures (Plate tectonic scale)

Continental regions adjacent to cool (upwelling) Western Sahara, Namibia (Africa) seawater zones. The amount of moisture evapo- Peruvian, Atacama, (South America) rated from the ocean surface is less under cooler Rub’Al Khali (Saudi Arabia) conditions, so yielding drier air. Cooler sea surface temperatures also stabilise air masses, reducing convective rainfall.

low pressure. As it rises it cools, losing most of the water vapour as rainfall to the tropical jungles and rainforests below. This now moisture-depleted air moves up and away from the equator and further cools and compresses. Finally, it sinks back to the earth‘s surface at around 30° latitude north and south of the equator. The cool descending dry air is reheated as it returns to the lower atmosphere, garnering an enhanced potential to absorb moisture, and so the major desert belts of the world are created (Fig. 5.1.5b). Local topography induces more localised variation in climate that can drive distribution and volume of evaporite sediments. For example, the playas and midlatitude deserts of central Asia and central Australia are in part a product of their isolation from the nearest ocean. Continentality is the term that describes the geographic isolation of a desert by its large distance from the ocean. It explains why the highly continental Takla Makan Desert of central China is a hyperarid desert at 50 °N latitude, well north of the world‘s subtropical high-pressure belts. The other factor helping form this type of desert is the rain shadow effect. Rain shadow deserts, also known as adiabatic deserts, form where air masses move up and over mountain ranges (Fig. 5.1.5c). As air rises over a mountain range it cools, moisture develops into clouds and falls as rain on the side of the mountain range facing into the prevailing wind. Subsiding air on the other side of the range is dry and its moisture-bearing capacity increases further as it sinks and warms. For example, the deserts of Patagonia lie on the lee side of the Andes, while the deserts of Nevada and

Utah are in the rain shadow of the Sierra Nevada. Adiabatic deserts can also be found behind the Himalayas, the Ethiopian uplands and Eastern Highlands of Australia. Extreme examples include the Dead Sea rift and the Afar depression; both areas are subsealevel depressions surrounded by high mountain ranges of the rift rim. Rain shadow effects increase where the region inland of the mountains lacks relief. Upwelling ocean waters off continental landmasses in the horse latitudes favour coastal deserts. Cold upwelling ocean water creates a belt of coast-parallel cool water. And, as evaporation is a function of heating, cold offshore oceans provide less moisture and less cloud to the adjacent land. The cool air above creates relatively dry onshore winds. As this colder air moves across the land it is heated, so developing an increased capacity to absorb rather than shed moisture (Fig. 5.1.5d). Offshore, these cool strips of water are frequently engulfed in winter sea fogs, created by a blanket of warmer air atop cold seawater. Such cold upwelling ocean currents account for the deserts of Baja California and the Namib of southwest Africa (Benguela Current). The Peru Current, as it upwells along the west coast of South America, helps maintain hyperaridity in the Atacama Desert of Chile, which receives around 1 mm of rain every 5-20 years making it the driest desert in today‘s world. Hadley cells are the broadest scale influence on today‘s desert occurrence and associated distribution of evaporite sediments. It is a planet-scale circulatory response to

Salt as sediment in the Central European Basin system as seen from a deep time perspective

Figure 5.1.6. Palaeolatitude distribution of evaporite deposits from Permian to present (in part after Ziegler et al. 2003; Warren 2006). This is a equal area plot whereby data plots are limited to one occurrence per rectangle measuring 5° of latitude and longitude. It does not indicate evaporite volume. The greater frequency of evaporites in the Northern Hemisphere simply reflects the greater preponderance of landmasses in the Northern Hemisphere in the time frame considered. Inset plots the latitudinal distribution of the area of landsurface covered by world’s modern deserts and also gives proportion of each continental area that is classified as arid or semiarid

varying intensity solar input from the equator to the poles (Table 5.1.1). Unfortunately, the climate of the Holocene is not well suited to be the baseline for the average Phanerozoic climate of the planet (Hay et al. 1997; Warren 2006). It is an interglacial period and, as indicated by the world δ18O curve from the deepsea sediments, is a state typical of only 10% of the icehouse climate of the past few million years. It is a time of relative sea-level stability after a rapid 130-m rise from the lowstand of the last glacial maximum. The last few million years is a period when physical geologic processes are operating at unusual rates and much of the earth‘s geochemical system is not in a steady state. During most of the Phanerozoic, except for part of the Permo-Carboniferous, there were no widespread continental or polar ice sheets and the planet‘s meridional temperature gradient was much less steep than it is at present. Yet for the past few million years, possibly since the Miocene and the initiation of permanent grounded Antarctic ice sheets, the distribution and effects of broadscale atmospheric circulatory belts (including Hadley cells) have been modified by the waxing and waning of

polar ice. During glacial periods the expanded polar ice sheets act as frigid plateaus, forcing a stronger meridional temperature gradient, and so alter significant portions of the world atmospheric circulation, as well as creating their own ice-cap associated pressure systems. Whenever icecaps expand, the Hadley-related circulation belts are pushed and compressed toward the equator. This has occurred numerous times during glacial maxima of the current icehouse climate mode. It has a marked effect on climate in Quaternary deserts, so that almost all haliteentraining continental playas have a sedimentary record of water-full versus dry stages in the last few hundred thousand years (e.g., Baker et al. 2001; Hesse et al. 2004; Yan et al. 1998). Times of expanded ice sheets, with their ice sheet-induced atmospheric circulation are probably the least representative states for the world‘s Phanerozoic climate. Interglacial periods within icehouse mode, like the Holocene, are perhaps a little closer to the “average” deep time climate. But today‘s climate is still a long way from a climate more representative of the bulk of Phanerozoic time when the polar regions were free of permanent ice and the world climate was in “greenhouse” not “icehouse” mode.

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In addition to glaciological intensity are another three factors influencing the broad scale distribution of the world‘s arid climate belts in deep time and their associated evaporite sediments. These factors are continentality, adiabatry and intensity of upwelling (Table 5.1.1) As we shall see, variations in the Phanerozoic scale of development of all three is controlled by the latitudinal position of continental plates and variations in the style of plate-plate development (i.e., the Wilson Cycle). When a supercontinent lies in an arid horse latitude belt, large portions of its interior will show a high degree on continentality. This was the case during the Carboniferous to Triassic in Central Europe, which at that time was part of the interior of the Pangaean supercontinent (Reinhardt and Ricken 2000). Likewise, adiabatic deserts are better developed at times when two continental plates collide. A rising mountain range is a barrier to atmospheric circulation, especially if perpendicular to the incoming circulation. For example, the Late Tertiary uplift of the Tibetan Plateau created an easterly jet stream, which now flows from Tibet across the Arabian Peninsula toward Somalia. This stream of dry air accentuated the existing aridity and created an adiabatic desert that lies almost on the Equator. In a similar fashion the Zechstein Basin lay in the continental interior of the Hercynean foreland. Adiabatic deserts and their evapor-

ite sediments will also form in arid climate zones when continental plates split. This is the case with the salts accumulating in the Danakil depression, which today lie in the rain shadow of the Danakil mountains and similar rain shadow effects would have influenced the various Mesozoic salt basins of the opening Atlantic. However, thick salt beds cannot form in any desert unless the arid basin depression can draw on a large brine reservoir. This was the case in the Miocene in the Red Sea area and the Aptian in the various proto-South Atlantic salt basins (Bosworth et al. 2005; Davison 2007). Both areas were rift-bound hydrographically-isolated subsealevel depressions able to draw on the huge reservoir of oceanic waters via seepage drawdown. Throughout the Phanerozoic such extensional sag and compressional plate interactions have helped induce hyperaridity atop large subsealevel seepage depressions that were capable of accumulating thick halite and gypsum successions (Warren 1999, 2006). Areas of meteoric-fed lacustrine halites that lack a marine source are an order of magnitude smaller over the same time frame (Warren 2006). Upwelling-induced deserts are also controlled by the positions of drifting continental plates within arid zones.

Figure 5.1.7. Volumes of Phanerozoic halite (after Hay et al. 2006). A) Mass of halite residing in basin. B) Reconstructed mass of recycled halite lost by combination of erosion and dissolution. Time of halite precipitation influenced by accretion and disintegration of Pangean supercontinent is shown

Salt as sediment in the Central European Basin system as seen from a deep time perspective

Today, the western edges of most horse latitude deserts are washed by cool ocean currents driven by large-scale ocean gyres, which rotate anticlockwise in the Southern Hemisphere and clockwise in the Northern Hemisphere. Throughout much of the Phanerozoic, coastal deserts in ancient horse latitude belts, as indicated by aeolian dunes and evaporites, have characterised the western sides of continents drifting into arid latitudinal belts (Hay 1996; Parrish 1982). When the latitudes of Holocene evaporites occurrences in modern deserts are plotted (as an equal area plot) against palaeolatitudes of ancient deserts, a bimodal latitude-related is obvious from the Permian to the present (Fig. 5.1.6; Ziegler at al. 2003). It indicates the time constancy for the broad distribution of the world‘s Hadley cells and the resultant horse latitude belt. This latitudinal similarity of evaporites through time is tied to planet-scale circulation in the world‘s atmosphere. As it is an equal area plot it does not indicate any time-related changes in the volume of salt accumulating within the world‘s two arid climatic belts located some 15-45° north and south of the equator. In fact, the volume of halite and other evaporite sediments does vary substantially with particular time periods (Fig. 5.1.7a; Hay et al. 2006)

5.1.4 Evaporites volumes in deep time There are periods in the past when evaporite deposition was more voluminous than today (Figs. 5.1.7, 5.1.8). The various evaporite basins, which together make up the Pangaean peak in the salt volume curve, were created mostly by plate tectonic mechanisms within what were tectonically induced arid to hyperarid climatic zones of the latest Palaeozoic to Mesozoic earth (Fig. 5.1.7a). Volumes of salts caught up in this association dwarf all Quaternary salt volumes and, as we shall see a little later, encompass a range of deposition styles for which we have very limited Quaternary analogues. Zharkov (1981) recognised 95 Permian evaporite basins with a total volume of 1.5 x 106 km3 of salts; 37 were halite-dominated, 48 were sulphate-dominated and 10 entrained substantial volumes of the potash salts. The majority of the evaporite sediments that make up the earlier part of the Pangaean salt peak first came into being as convergent basins that formed when the continents collided in the Carboniferous and Permian and so constructed the Pangaean supercontinent. Carboniferous to early Permian megahalites were precipitated in the Williston Basin and northern Ellesmere

region in North America, in northern Brazil, North Africa, the East-European Platform, and Kazakhstan (Fig. 5.1.8). These Carboniferous to early Permian collision basins and farfield sags define a broad belt along the western Pangaea margin, but also in northeastern North America and eastern Greenland, as well as in northern South America, including western Brazil. Large salt deposits of the same age are found in central Europe including portions of the East-European Platform and the Urals. The largest Early Permian deposit fills the East European Basin from the Caspian to the Black Sea with some 500 x 103 km3 of halite and 100 x 103 km3 of anhydrite (Magaritz 1987). Subsequent Late Permian intracontinental salts include major deposits in the Zechstein Basin of northwestern Europe and the evaporites of west Texas (Fig. 5.1.1a). The Zechstein deposits were not solely a farfield intracratonic farfield sag response, but include an extensional component in their tectonic genesis (see later). Together all these basins define a broad band of equatorial aridity extending from about 50 °N to about 30 °S palaeolatitude (Fig. 5.1.6), with basinwide salt deposition ranging in time and world climate across Carboniferous icehouse to Permian greenhouse (Fig. 5.1.7). The even larger post-Permian Pangaean salt peaks (Late Triassic-Cretaceous) indicate times when Pangea rifted and the divergent basins of the opening Atlantic and Tethys passed into hydrographically-isolated subsealevel positions (Fig. 5.1.8). During the Early Triassic, the amount of evaporite precipitation intensified as huge intracontinental and continental margin halite pans associated with redbeds developed in southeastern and western USA, western Canada, northwestern Africa, western and southeastern Europe, central South America, south-central Africa and eastern Greenland (Fig. 5.1.7). With the passage into a greenhouse climate mode, large sulphate platforms were established in northern Arabia, Iraq, Kazakhstan, and northern China. This increased salt volume from the Triassic to the Cretaceous in large part reflects intensified rift activity in the horse latitudes during the break-up of Pangaea. Extension, followed by cooling and sagging, established a number of large isolated subsealevel rifts and intracratonic sags, which acted as the preeminent salt accumulators. Thus, the Pangaean evaporite peak ties to three timerelated plate tectonic associations (Table 5.1.2): (i) Pennsylvanian to Permian accretion of the Pangaean supercontinent, where marine-fed collision basin accumulated thick salt beds along the irregularly embayed margins of Gondwana, Laurussia and Siberia; (ii) Mesozoic intraplate extensional depressions created during the subsequent break-up of Pangaea into Gondwana/Laurasia; salt beds deposited in these widening marine-fed rift basins, including the continental salts of the Newark Basin (Tri-

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Figure 5.1.8. Distribution of halite-entraining basins across deep time (figure is compiled with permission, as a GIS output in a Mercator projection from a commercial evaporite database, courtesy of J K Resources, Pty Ltd.). This plot does not illustrate calcium sulphate-dominant systems. Basins: 1. Abenaki (N. Scotian); 2. Adavale; 3. Adelaide fold belt; 4. Adriatic -Albanian foredeep; 5. Amadeus Basin; 6. Amadeus Basin (Chandler); 7. Amazonas; 8. Andean; 9. Andean; 10. Apennine; 11. Appalachian; 12. Aquitaine; 13. Arabian Basin (Gotnia Salt Basin); 14. Arabian Basin (Hith Salt Basin); 15. Arabian Basin (Hormuz central Saudi); 16. Arabian Basin (Hormuz Gulf region); 17. Arabian Basin (Hormuz-Kerman region); 18. Atlas (Algerian - Tunisian); 19. Atlas (Moroccan); 20. Baltimore Canyon; 21. Berrechid; 22. Betic-Guadalquivir Basin; 23. Bohai Basin; 24. Bonaparte (Petrel); 25. Brazilian Aptian Basin (Camamu); 26. Brazilian Aptian Basin (Campos-Santos); 27. Brazilian Aptian Basin (Ceara); 28. Brazilian Aptian Basin (Cumuruxatiba); 29. Brazilian Aptian Basin (Sergipe-Alagoas); 30. Cankiri-Corum; 31. Canning Basin; 32. Cantabrian-West Pyrenees; 33. Carnarvon Basin (Yaringa); 34. Carpathian foredeep; 35. Carson Basin (Grand Banks); 36. Chu-Sarysu (Devonian); 37. Chu-Sarysu (Permian); 38. Cicilia-Latakia; 39. Cuban; 40. Danakil; 41. Dead Sea; 42. Dniepr-Donets; 43. Dniepr-Donets; 44. Eastern Alps; 45. Ebro Basin; 46. Flemish Pass Basin (Grand Banks); 47. Georges Bank; 48. Green River Basin; 49. Gulf of Mexico (Northern; Gulf coast); 50. Gulf of Mexico (Southern; Salina-Sigsbee); 51. Haltenbanken; 52. Haymana-Polatli; 53. Holbrook Basin; 54. Horseshoe Basin (Grand Banks); 55. Hudson Bay; 56. Ionian; 57. Jeanne d’Arc Basin (Grand Banks); 58. Jianghan Basin; 59. Jura/Rhodanian; 60. Katangan; 61. Khorat Basin; 62. Kuqa Foreland (Tarim Basin); 63. La Popa (Monterrey) Basin; 64. Lusitanian; 65. Mackenzie Basin; 66. Maestrat; 67. Majunga Basin; 68. Mandawa Basin; 69. Ma’Rib-Al Jawf/Shabwah (Hadramaut); 70. Maritimes Basin; 71. Mediterranean–Western; 72. Mediterranean–Adriatic; 73. Mediterranean-Andros Basin; 74. Mediterranean-Cretean Basin; 75. Mediterranean - Samothraki basin; 76. Mediterranean-Tyrrhenian; 77. Mediterranean-Central; 78. Mediterranean-Eastern; 79. Mediterranean-Sicilian; 80. Michigan Basin; 81. Moesian; 82. Moose River Basin; 83. Neuquen Basin; 84. Nordkapp Basin; 85. Officer Basin; 86. Olduvai depression; 87. Oman (Fahud Salt Basin); 88. Oman (Ghaba Salt Basin). 89. Oman (Ghudun Salt Basin); 90. Oman (South Oman Salt Basin); 91. Oriente-Ucayali (Pucara) Basin; 92. Orpheus Graben; 93. Palmyra; 94. Paradox Basin; 95. Parry Islands Fold Belt; 96. Pricaspian Basin; 97. Pripyat Basin; 98. Qaidam Basin; 99. Qom-Kalut; 100. Red Sea (North); 101. Red Sea (South); 102. Rot Salt Basin; 103. Ruvuma Basin; 104. Sabinas Basin; 105. Sachun Basin; 106. Salar Basin (Grand Banks); 107. Salt Range (Hormuz - Punjab region); 108. Salt Range (Kohat Plateau); 109. Saltville (Appalachian); 110. Scotian Basin; 111. Siberia, East; 112. Sirjan Trough; 113. Solimoes; 114. Somalia-Kenya; 115. South Whale Basin (Grand Banks); 116. Sverdrup Basin (Ellef Ringnes - NW Ellesmere); 117. Sverdrup Basin (Melville Is); 118. Tabriz Salt Basin; 119. Tadjik Basin; 120. Takutu Salt Basin; 121. Transylvanian; 122. Tromso Basin; 123. USA Midcontinent; 124. West Africa (Angola-Gabon); 125. West Africa (Gambia-Guine Bissau); 126. West Africa (Mauritania-Senegal); 127. West Africa (Morocco-S. Spain); 128. Western Canada (Alberta Basin); 129. Whale Basin (Grand Banks); 130. Williston Basin; 131. Zagros (Mesopotamian Basin); 132. Zagros (Mesopotamian Basin); 133. Zechstein (NW Europe); 134. Zechstein (onshore UK); 135. Zipaquira Basin

Salt as sediment in the Central European Basin system as seen from a deep time perspective

assic), the marine salts of the Canadian offshore (Triassic), the Gulf of Mexico (Jurassic), the Campos Basin (Cretaceous) and the various mostly marine-fed salt basins of offshore west Africa (Triassic-Jurassic in central and northern region, Cretaceous in Gabon-Angola salt basins; and (iii) Associated intracratonic and intracontinental Mesozoic sag salt accumulations in basins that typically sit atop the failed arms of earlier more sharply-defined fault-bound rift successions or lie in the hinterland or foreland of orogenic belts created by farfield stresses driven by interplate collision (Fig. 5.1.8). The intracratonic-sag basins in the Pangaean association are especially obvious in the Permian Zechstein and various Triassic and Jurassic salt basins of onshore Northern Africa, which formed isolated subsealevel marine-fed depressions as Pangaea unzipped along its Tethyan and Atlantic axes. Large intracratonic sags also typify many other saline giants outside of the time of the Pangaean supercontinent (e.g., the Devonian Elk Point Basin, the Palaeozoic Williston Basin, the Silurian Michigan Basin and Ordovician Canning Basin), as do other rift (e.g., Frasnian of Dniepr-Donets of eastern Europe) and collision (Miocene of the Alpine Mediterranean and the Zagros foreland) associations (Fig. 5.1.8). Creation and destruction of the Gondwana supercontinent in the Neoproterozoic-Early Cambrian (≈550Ma; a.k.a. Infracambrian) defines an earlier, but not as well documented, peak in the halite volume curve shown in figure 5.1.7a. Thick layers of halokinetic rock salt (halite) are widespread in the latest Ediacarian to the Early Cambrian centred on an area that now lies largely in the Middle East (Fig. 5.1.8). Indeed, the Hormuz strata contain some of the world‘s thickest successions of rock salt (Edgell 1996). They include the Hormuz Salt of Iran and the Arabian Gulf, the Ara Salt of Oman (both units are thought to have begun to accumulate in be latest Ediacarian and to straddle the Cambrian boundary), the Salt Range salt of Pakistan (Atdabanian-Botoman), and the Usolka and equivalent salts of Siberia (TommotianAtdabanian). The assembly of Gondwana began during the Ediacarian to Early Cambrian. It involved the amalgamation of the separate crustal blocks of Avalonia, Europa, Arabia, Africa, Madagascar, South America, and Antarctica (together forming West Gondwana) and resulted in the compressional Pan-African orogeny, which culminated between 560 and 530 Ma. East Gondwana (India, South China, North China and Australia) collided with West Gondwana along the Mozambique suture between ca. 600 and 550 Ma. The development of thick salts in the Ara Formation, once thought to be rift deposits of Tommotian age (Husseini and Husseini 1990), now appear to be foreland basin deposits of late Ediacarian age (Al-Siyabi 2005).

There is evidence for an even older aridity-cored supercontinent (Neoproterozoic ? 800 Ma) in what is now largely the interior of the Australian continent (Fig. 5.1.8). Remnants of these supercontinent evaporites reside in the thick halite beds of the Bitter Springs Fm. in the Amadeus Basin and the halites of the Browne and Madley formations in the Officer Basin of Central Australia, as well as the widespread dissolution breccia (rauhwacke) after Callana Fm. halites in the Flinders Ranges. The rauhwacke of the Duruchaus Fm. in Namibia and its equivalents in the Katangan copper Belt are perhaps also related to the break-up of the Mesoproterozoic supercontinent, known as Rodinia or Kanatia, some 800 Ma. The great age of these deposits and their subsequent polyphase tectonic history makes problematic the exact determination or their tectonic situation at the time the thick halite beds were laid down. They are typically classified as intracratonic, but they are also aulocogens, so that they are also considered by some as synrift or early postrift deposits tied to the break-up of Rodinia (Li and Powell 2001; Veevers 2004). Volumes of evaporites deposited in many of these ancient saline giants are huge, with individual deposits accumulating over geologically short time frames of no more than a few million years. For example, the deposition of the Jurassic Louann Salt in the Gulf of Mexico, some 175 Ma, sequestered approximately 8% of world-ocean NaCl (Land et al. 1988), while the Messinian evaporites of the Mediterranean were deposited in less than 300,000 years yet lowered oceanic salinities by some 2-4‰. The extraction of the large volumes of NaCl from the Mesozoic oceans probably explains the 12‰ drop in seawater salinTable 5.1.2. Timing of the assembly and disassembly of the Pangean supercontinent (after Warren 2006). Late Palaeozoic formation of Pangaean supercontinent

Pennsylvanian collision of Laurussia, Kazakhstania and Siberia to form Laurasia (≈280 ma), followed by Permian collision of Laurasia and Gondwana

Mesozoic breakup of Pangean supercontinent

Breakup of Pangaea: 220-190 Ma: Opening of North Atlantic Newark and other rift basins in NE North America and NW Africa 175 Ma: Opening of Neo-Tethys 155 Ma: Opening of Palaeo-Gulf of Mexico 125 Ma: Separation of South America from Africa (Aptian opening of south Atlantic)

Associated Late Carboniferous orogenic events: Appalachian/Alleghenian orogeny of E North America Mauritanide/Moroccan fold belt of NW Africa Hercynian (Variscan) orogeny of S Europe Marathon/Ouachita orogeny of S North America Uralian orogeny of Europe-Siberia

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J.K. Warren Figure 5.1.9. Similar-scale plot comparing aerial extent of Permian Zechstein and Triassic Röt salt basins of Europe with various Quaternary regions of the world’s deserts where bedded evaporite sediments are precipitating (figure is compiled with permission, as a GIS output in a Mercator projection from a commercial evaporite database, courtesy of J K Resources, Pty Ltd.)

ity noted by Hay et al. (2006) across this time frame (Fig. 5.1.4c). But halite is a highly soluble salt and, as well as being lost from the oceans via deposition, these same evaporite sediments are susceptible to recycling their solutes back into the world‘s oceans by ongoing erosion and dissolution, especially when a thick halite succession becomes halokinetic, as in the various circum-Atlantic salt basins. When the volume of salt being recycled into the world’s oceans is accounted for, as was done by Hay et al. (2006) an indication of actual volumes of salt deposited in a megahalite basin can be reconstructed via an accounting of the volume of salt still in the basin and the likely volume recycled (Hay et al. 2006). The resulting Phanerozoic reconstruction (Fig. 5.1.7b) shows how our notions of salt volume through time are strongly influenced by the volume of salt still preserved in the basin. The largest bulge in the volume of salt curve resides in the salt held in Jurassic and Cretaceous basins and relates to the unzipping of Pangaea. But in a recycled salt reconstruction it seems that there were similarly large volumes of salt deposited on the Pangaean supercontinent as its constituent plates came together in the Late Carboniferous (Fig. 5.1.7b). The reconstructed curve also implies that the largest volume of salt ever deposited in a sedimentary system in the last 700 million years did so in the Infracambrian basins of the Middle East. We have no Quaternary counterpart to the volumes and aerial extents of evaporite sediments that accumulated in these ancient Phanerozoic and Neoproterozoic basins. As an example, consider figure 5.1.9, it is a same scale, sameprojection plot of the Permian Zechstein and Triassic Röt evaporite basins of Europe overlain by the aerial extent of the largest Quaternary examples of regions where bed-

ded evaporites accumulate today. Clearly, there is a scale problem comparing modern and ancient salt deposits. So let us now attempt to answer why Quaternary analogues offer such a limited solution to the puzzle of scale and thickness in ancient salt sediments.

5.1.5 Evaporite volumes & tectonics? For a thick halite bed to accumulate requires a permanent halite-saturated brine curtain that bathes the halite bed as it subsides and is buried. These brine curtains require a combination of subsidence and ongoing brine reflux to maintain a longterm position in earth space (Warren 2006). Of course, “longterm” is a relative term. By “longterm” I mean that a halite-saturated brine curtain is present throughout the time the halite bed or beds accumulate. When the hydrology is suitable, as it was for a few hundred thousand years in the Late Miocene (Messinian) in the Mediterranean, then an evaporite bed more than a kilometre thick could accumulate in the lower parts of hydrographically isolated marine basin floor (Fig. 5.1.8) Today, and throughout the Quaternary, evaporite beds up to a few hundred metres thick accumulate only in continental hydrologies in arid to hyperarid settings, exemplified by the salars of South America such as Salar de Atacama and Salar de Uyuni, the Dead Sea in the Middle East and perhaps Death Valley in California (Fig. 5.1.9). All such hydrologies typify the bottoms of relatively steep-sided endoheic drainage basins and are characterised by basin floor subsidence tied to an active tectonic setting. This combination of factors creates accommodation space for the accumulation of saturated continental brines in the lowest part of the subsiding basin and facil-

Figure 5.1.10. Tectonic associations where thick, bedded sequences of clean pure halite can accumulate (after Warren 2006)

Salt as sediment in the Central European Basin system as seen from a deep time perspective 263

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Table 5.1.3. Listing of tectonic setting of various halite-entraining basins at the time that thick halite was precipitated. It is the synprecipitational setting, so and in basins with an ongoing polyphase tectonic history (like the various Zechstein sub-basins of Europe) it is not the same as the salt’s current tectonic situation (data output from GIS database courtesy of J K Resources Pty. Ltd). See figure 5.1.8 for basin age and location. Asterisk indicates postrift megahalite deposition. Divergent

Abenaki (N. Scotian), Adelaide fold belt, Adriatic -Albanian foredeep, Apennine, Aquitaine, Arabian Basin (Gotnia Salt Basin), Arabian Basin (Hith Salt Basin), Atlas (Algerian-Tunisian, Atlas (Moroccan), Baltimore Canyon, Berrechid, Betic-Guadalquivir Basin, Brazilian Aptian Basin (Camamu), *Brazilian Aptian Basin (Campos-Santos), Brazilian Aptian Basin (Ceara), Brazilian Aptian Basin (Cumuruxatiba), *Brazilian Aptian Basin (Sergipe-Alagoas), CantabrianWest Pyrenees, Carson Basin (Grand Banks), Chu-Sarysu (Devonian), Cuban, Danakil, Dniepr-Donets, Eastern Alps, Flemish Pass Basin (Grand Banks), Georges Bank, Gulf of Mexico (Northern; Gulf coast), *Gulf of Mexico (Southern; Salina-Sigsbee), Haltenbanken, Horseshoe Basin (Grand Banks), Ionian, Jeanne d’Arc Basin (Grand Banks), Jura/ Rhodanian, Katangan, La Popa (Monterrey) Basin, Lusitanian, Maestrat, Majunga Basin, Mandawa Basin, Ma’RibAl Jawf/Shabwah (Hadramaut), Moesian, Nordkapp Basin, Olduvai depression, Orpheus Graben, Palmyra, Red Sea (North), Red Sea (South), Rot Salt Basin, Ruvuma Basin, Sabinas Basin, Salar Basin (Grand Banks), Scotian Basin, Somalia-Kenya, South Whale Basin (Grand Banks), Tadjik Basin, Takutu Salt Basin, Tromso Basin, *West Africa (Angola-Gabon), West Africa (Gambia-Guine Bissau, West Africa (Mauritania-Senegal), West Africa (Morocco-S. Spain), Whale Basin (Grand Banks), Zechstein (NW Europe), Zechstein (onshore UK)

Convergent

Adavale, Amadeus Basin (Chandler), Andean, Arabian Basin (Hormuz central Saudi), Arabian Basin (Hormuz Gulf region), Arabian Basin (Hormuz-Kerman region) Bohai Basin, Cankiri-Corum, Carpathian foredeep, Chu-Sarysu (Permian), Cicilia-Latakia, Ebro Basin, Green River Basin, Haymana-Polatli, Holbrook Basin, Jianghan Basin, Khorat Basin, Kuqa Foreland (Tarim Basin), Mediterannean-Western, Mediterranean-Adriatic, Mediterranean-Andros Basin, Mediterranean-Cretean Basin, Mediterranean-Samothraki Basin, Mediterranean-Tyrrhenian, Mediterranean-Central, Mediterranean-Eastern, Mediterranean- Sicilian, Neuquen Basin, Oman (Fahud Salt Basin), Oman (Ghaba Salt Basin), Oman (Ghudun Salt Basin), Oman (South Oman Salt Basin), Paradox Basin, Pricaspian Basin, Qaidam Basin, QomKalut, Sachun Basin, Salt Range (Hormuz - Punjab region), Salt Range (Kohat Plateau), Sirjan Trough, Sverdrup Basin (Ellef Ringnes - NW Ellesmere), Sverdrup Basin (Melville Is), Tabriz Salt Basin, Transylvanian, Western Canada (Alberta Basin), Zagros, Zagros (Mesopotamian Basin), Zipaquira Basin

Intracratonic

Amadeus Basin, Amazonas, Appalachian, Bonaparte (Petrel), Canning Basin, Carnarvon Basin (Yaringa), Hudson Bay, Mackenzie Basin, Maritimes Basin, Michigan Basin, Moose River Basin, Officer Basin, Oriente-Ucayali (Pucara) Basin, Parry Islands Fold Belt, Pripyat Basin, Saltville (Appalachian), Siberia-East, Solimoes, USA Midcontinent, Williston Basin

Transform

Dead Sea

itates the focusing of hypersaline waters into this deepest part of the basin, so allowing the maintenance of a brinesaturated hydrology where salt beds that are tens-of-metres-thick can stack. However, all these Quaternary examples of stacked relatively pure halite beds are limited by the volume of water that can be supplied to a basin. Hydrogeochemical and isotopic studies of waters in these Quaternary systems show that the majority of solutes needed to accumulate the lacustrine halite do not come from direct rainfall or surface runoff (e.g., Rosenthal et al. 2006; Martini 1969; Hardie 1991). Rather the solute source is brackish to saline groundwaters leaching solutes as they seep into the depression via the surrounding aquifers. Thus the volume of appropriate brine that can be created in such a meteoric hydrology limits the nonmarine system. If the volume of groundwater flowing into the depression increases, as it has during more humid Pleistocene (glacial) episodes then the salinity of the system decreases and lacustrine mudstones are deposited, typically on the bottom of a perennial, somewhat fresher, lake. If on the other hand

the degree of aridity increases then the volume of water entering the system decreases. Self-limiting brine volume inflows in tectonically active continental settings, and the ability of artesian meteoric waters to destroy a longterm brine curtain in less tectonically active settings, is why modern and ancient lacustrine evaporite deposits tend to lie at the smaller scale end of our spectrum of areal extents of ancient halite-entraining basins (Fig. 5.1.9). Hydrologies that facilitate deposition of thick sulphate or halite evaporites across large parts of a basin floor are related to settings that allow huge volumes of seawater to be drawn into a depression and then concentrate. Because water cannot flow uphill, all systems drawing on seawater as the primary source of solutes must by definition be subsealevel or drawdown hydrologies. Hydrological conditions must be characterised by hydrographic isolation (no permanent connection to the ocean) and seepage of seawater must be the main hydrological mechanism supplying brine (see Warren 2006; Chapters 2 & 5 for more detail and the relevant references to this hydrological argument).

Salt as sediment in the Central European Basin system as seen from a deep time perspective

Older buried salt masses (especially if halokinetic and near the surface) can also contribute solutes to continental drainage inflows in endoheic depressions. This is the case in the halite-entraining salars of South America, the Holocene potash deposits of interior China, and was probably the case when the Miocene Fars/Gachsaran salts were deposited in the Zagros foreland. There are only a few modern systems where drawdown is allowing the accumulation of marine-seepage derived halite beds. They include the gypsum/halite coastal salinas of southern and Western Australia and Lake Asal in the Red Sea region (Fig. 5.1.9; Logan 1987; Warren 1982a,b; Gasse and Fontes 1989). All are small scale and are not indicative of the scale and volume of salts possible when appropriate seawater drawdown hydrologies have set up in the past in appropriate tectonic and climatic associations. Clearly, ancient large evaporite basins are the result of a combination of tectonic and hydrological circumstances that are not currently active. They were last active in the Miocene in association with collision basins tied to the Alpine-Himalaya orogenic belt and to rift basins developed in the early stages of the Red Sea. They will be active again in the future at sites and times of appropriate plate-plate interaction when two continental plates are in close proximity. Continent-continent proximity allows the floor of the intervening depression to be subsealevel and large, but hydrographically isolated (Warren 2006). Timing and styles of plate-plate interaction are summarised by the Wilson Cycle, which breaks down a plate tectonic cycle into a time of divergence passing to convergence via stages A through H (as shown in figure 5.1.10). Named in honour of J. Tuzo Wilson, the Wilson Cycle encapsulates the notion that as new ocean basins open, via the formation of new seafloor, others must close driven by the process of subduction and ultimately by continentcontinent collision (Wilson 1966). There are four plate settings where thick evaporite beds have accumulated at particular times in the Phanerozoic and now defined by the obvious peaks in figure 5.1.7. They are (Fig. 5.1.10; Table 5.1.3): (i) Areas of rifting where continental plates are beginning to move apart (divergent plates in rift to drift transition in Wilson cycle stages B to early C); (ii) Areas where continental plates are coming close to one another (convergent or colliding plates in Wilson cycle stage F); (iii) Areas of intracontinental or intraplate sagging, which are often the result of farfield plate stress induced by continental orogeny acting on zones of pre-existing crustal weakness such as an old rift belt (aulocogen) or farfield sags tied to the formation of foreland depressions (intracratonic basins in Wilson cycle stage A). (iv)

Thick (km-scale) halite sequences in regions of rapidly subsiding continental crust in transform or strike-slip settings where the plates are sliding horizontally past one-another (conservative or neutral plate margin). Only in Wilson stages F (continent-continent collision) through the onset of Stage C (intracratonic through rift through incipient ocean) can continental plates be in sufficient proximity to create the hydrographic isolation needed to facilitate marine seepage to a desiccated subsealevel basin floor and so precipitate marine-fed saline giants. Stages C, D and E encapsulate open marine situations atop ocean crust with a hypsometry that requires complete hydrographic connection to the ocean for all basins, so precluding any possibility of subsealevel hydrographic isolation. Brine inflow to the various tectonically-induced evaporitic lows tends to be maintained by continuous seepage of marine groundwaters into the depression (i.e., they are subsealevel marine-fed hydrologies). This ancient evaporite hydrology is always tied to an arid setting as these are the zones of the world‘s surface where there is always the potential to lose more water from the basin than can actually enter via a combination of groundwater seepage and surface floods. And so, throughout deep time, evaporite sediment occurrence characterises the same two latitudinal belts that characterise Holocene evaporites (Fig. 5.1.6). But, what is not constant is the volume of evaporite sediment deposited in the horse latitudes (Fig. 5.1.7). As we have seen latitudinal similarity of evaporite distribution through time is tied to the Hadley circulation in the worlds atmosphere, but latitude does not indicate volume of salt. For large volumes of marine-fed salt to accumulate in and around ancient horse latitude belts also required the appropriate combination of tectonics and climate in zones of continent-continent plate interaction. This is why there are no Quaternary examples of megahalites. Our largest and thickest bedded halites are forming in suprasealevel Quaternary collision belts as lacustrine “piggy-back” basins in the salars of South America (Bobst et al. 2001; Jordan et al. 2007). The only Quaternary examples of marine-fed subsealevel rifts are smallscale and include Lake Asal, Djibouti and the Pleistocene marine-fed halites (with MgSO4 bitterns) in the Danakil depression (Gasse and Fontes 1989; Hutchinson and Engels 1970). The Holocene inflow waters of the Danakil are continental, not marine and the bitterns lack MgSO4. In the recent geological past the best example of a full-scale rift-associated marine seepage evaporite is probably the thick Miocene salt in the Red Sea (Bosworth et al. 2005). Other older examples of rift-associated evaporites include the circum-Atlantic salt basins (Fig. 5.1.8; Manspeizer 1982).

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J.K. Warren Figure 5.1.11. Stages in rift related sedimentation showing typical characteristics of prerift, synrift and postrift sediment packages

Evaporites deposited in ancient continental rifts that passed into incipient ocean basins have a sediment fill that can be further subdivided in prerift, synrift and postrift (Fig. 5.1.11). Prerift accumulations are lithologies that predate the onset of the extensional faulting that defines the rift. Synrift sediments are deposited in a subsiding zone that thickens toward active normal fault(s) along the rift margin. There is typically a well-defined rift onset unconformity at the lower contact of the synrift with the prerift sediments. The synrift sediment package typically expands over an ever-larger area with time driven by ongoing fault activity. Expansion in the area of synrift sediments is associated with onlap of sedimentation surfaces and is a characteristic feature that is usually recognisable in associated seismic reflectors. Post-rift sediments are deposited when the rift-bounding faults are no longer active. These sediments tend to be deposited atop and over the rift graben edges. When this classification of pre-rift, synrift and postrift sediment packages is applied to thick salts that accumulated at divergent plate margins it seems most megahalites in divergent settings were deposited as synrift packages within well-defined graben profiles exemplified by the various Triassic salt basins of the northern Atlantic rim. Permian and Triassic salt basins of Europe and the North Sea are also synrift deposits (Fig. 5.1.8, Table 5.1.3; Hudec and Jackson 2007; Warren 2006). The significance of the synrift style of deposition on subsequent halokinesis in the European region, which is an area with a polyphase tectonic history and multiple salt layers in the post Carboniferous stratigraphy, is discussed later in the chapter. A plot of world salt basins at divergent margins shows the largest and thickest halite accumulations are postrift, not synrift packages (Fig. 5.1.8, Table 5.1.3). They formed in palaeo-Gulf of Mexico and in the South Atlantic Aptian salt basins of Brazil and West Africa (Fig. 5.1.8). The lateral extent into the synprecipitational basin of these deposits were not defined by a graben edge atop transitional oceanic crust, but in places thick postrift halites were in direct contact with seafloor basalts of

the oceanic crust. It is very likely that the hydrographic barrier that facilitated the hydrology of these subsealevel megahalite basins was also an offset volcanic ridge. It also means that the northern and southern Jurassic salt masses of the Proto-Gulf of Mexico, and the Brazilian and Angolan Aptian salt masses were never in direct depositional contact, but were separated by a basaltic midoceanic ridge, even as they formed (Salvador 1987; Jackson et al. 2000). Recently, Huismans and Beaumont (2008) proposed that, compared with earlier one dimensional depth dependant models, the postrift Mesozoic salt basins of the Atlantic are better explained by a model that includes notions of decoupling between upper and lower parts of the lithosphere during stretching, contrasting wide and narrow extensional styles above and below the decoupling level, and progressive focusing of crustal extension toward the rift axis. Halokinesis in postrift evaporite-floored basins is characterised by well-developed allochthonous salt tiers,

Figure 5.1.12. Lateral extents of megahalite basins in synrift and post rift settings

Salt as sediment in the Central European Basin system as seen from a deep time perspective Figure 5.1.13. Zechstein Salt. A) Salt structures in the Zechstein Basin of the Souther North Sea and NW Europe showing halokinesis is best developed in the more central parts of the Zechstein Basin (after Coward and Stewart 1995) B) Facies distribution in the southern Zechstein Basin (after Geluk 2000). The basin is bordered to the north by the Mid North Sea High and RingkøbingFyn High and the London-Brabant Massif and Rhenish Massif in the south. Platform is rich in carbonate, basin is halite-rich. Note the location of the Variscan Thrust Front. MNSH: Mid North Sea High; RFH: RingkøbingFyn High; SB: Silverpit Basin, LBM: London-Brabant Massif; RM: Rhenish Massif; NGB: North German Basin; PT: Polish Trough. Dashed rectangle indicates area encompassed by figure 5.1.13a

strong components of subhorizontal salt translation and secondary minibasin-capped canopy levels (Aptian salt basins of the South Atlantic and the Gulf of Mexico; Davison 2007; Tari et al. 2003; Hudec and Jackson 2006). Stronger vertical aspects of salt flow (cf. lateral flow), both in extension and compression, characterise synrift basins such as the Permian Zechstein Basins of Europe and the Triassic salt basins of the circum North Atlantic (Fig. 5.1.8, Table 5.1.3).

5.1.6 Episodic halokinesis The flat-bedded nature of much of the halite in the Zechstein basin of central Europe has been modified to varying degrees by the basin’s subsequent tectonic history. Flowing salt in the Central European Basin system reached the desert landsurface as namakiers (salt glaciers) during extensional episodes in the Triassic. Bedded salt was being simultaneously deposited in the lower parts of the

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variety of salt pillows and diapirs that create subsalt and suprasalt hydrocarbon traps (Figs. 5.1.13a, b). Ridges and walls of Zechstein salt show a regularity of ridge spacing and orientation tied to the major fault trends that were active at the time the salt was flowing with many diapirs underlain by basement faults (Coward and Stewart 1995). During episodes of Triassic extension, underlying presalt ridges created barriers to ongoing salt flow and so focused the movement of salt in their vicinity (Fig. 5.1.14). Many extension-related Triassic and Jurassic ridges were reactivated, buckled and shortened by subsequent Late Cretaceous compression and basin inversion (Mazur et al. 2005; Remmelts 1995; Davison et al. 2000, Davison 2007).

Figure 5.1.14. Dominant halokinetic styles in structures composed of Zechstein salt during Triassic extension and Late Cretaceous compression

same basin. In this section we will explore depositional analogues and possible feedbacks between namakiers and the deposition of bedded Triassic halite. Salt flows fed by Zechstein salt also reached the seafloor during extensional episodes in Jurassic history of the North Sea and again during shortening episodes on the Late Cretaceous chalky seafloor. But the focus of this chapter is bedded salt, not a general discussion of halokinesis, so we focus on namakiers from a deep time perspective and their possible influence, via ongoing dissolution and reworking, on the formation of contemporary bedded salt. We will not attempt an overview of the diversity of salt structures that occur in the Central European Basin or the North Sea. The interested reader is referred to Stewart (2007) for an excellent overview of this topic and more general aspects of halokinesis are discussed in Jackson et al. (1995); Alsop et al. (1996) and Warren (2006). Today, most halokinetic structures sourced in the Zechstein halite are in more central basin positions, where the salt is purer and thicker. There the salt is caught up in a

That Zechstein salt did not flow until tens of millions of years after the accumulation of salt had ceased, is evidenced by the constant prekinematic thickness of much of the Buntsandstein of the Triassic section of NW Germany (Fig. 5.1.15). Regionally, much of the irregular thickness of the Triassic section in areas west and east of the Central North Sea is due to slip of the Triassic carapace above Zechstein salt, leading to reactive diapirism of salt between Triassic thin skinned rifts, or to pillow-like buckle folding of the Trias above the salt (Coward 1995). This salt movement produced a series of essentially north south-trending salt rises separated by fault-defined extensional depressions that were the sites for thicker Triassic deposition (Fig. 5.1.13a). Episodes of deposition of bedded salt within the various arid Triassic grabens of central Europe correspond to times of basin extension and substantial Zechstein salt movement (Fig. 5.1.2a; Mohr et al. 2005; Mazur et al. 2005). This style of salt flow, responding to regional tectonic extension some tens of millions of years after thick Zechstein salt was first deposited, contrasts with the nature of salt flow in the postrift basins of the Gulf of Mexico and the Aptian salt basins on both sides of the South Atlantic. For example, the Jurassic Louann salt-cored structures and associated salt glaciers began to flow during, or very soon after, the deposition of beds immediately atop the halite, especially in the updip portions of the palaeo-Gulf of Mexico (e.g., Jurassic Norphlett and Smackover For-

Figure 5.1.15. Simple tripartite breakdown of deposition of Zechstein salt movement (as related to Mesozoic extension) into prekinematic (Lower to Middle Buntsandstein, synkinematic (Upper Buntsandstein, Muschelkalk, Lower to Middle Keuper) and post kinematic layers. For location and detailed stratigraphic breakdown of this onshore E-W seismic line in NW Germany see figures 3 and 4 in Mohr et al. (2005; seismic image courtesy of Marcus Mohr)

Salt as sediment in the Central European Basin system as seen from a deep time perspective

Figure 5.1.16. Triassic salt flow in Keuper time in NW Germany (after Mohr et al. 2007). A) Seismic line from NW Germany showing typical Christmas tree profile with Triassic namakier tongues extending away from salt stem onto surrounding sediments. B) Variance map looking down on a salt cored anticline at a level 50 m below the unconformity at base of the Cretaceous displays a discontinuity in the western part of the image due to a fault-bound fan-like namakier (both seismic images courtesy Marcus Mohr)

mations; Kopaska-Merkel and Mann 1994; Fleming and Meyland 2000). That salt did not flow so early in burial history of the Zechstein basin is probably a response to the largely continental nature of the crust beneath the Zechstein salt basin. This has led some to classify the 150-300 m thick (at deposition) Zechstein salt as being deposited in an intracratonic basin (e.g., Hudec and Jackson 2007). The thin bedded (<50 m thick) intercalated nature of the various overlying Triassic salt beds is more typical of an intracratonic salt deposit (Warren 2006). In my opinion, the faulted graben-controlled nature of sediment beneath Zechstein salt, the fact that Zechstein platform reefs and other associated carbonate facies belts show a syndepositional relief tied to ongoing extensional faulting (e.g., Glennie 1986; Geluk 1999a; Ziegler 1989), makes the Zechstein salt association synrift rather than intracratonic for onshore Europe (Table 5.1.3). I think of the Röt halite being postrift and a more typical intracratonic deposit. However, the classification and the subsidence mechanism creating the accommodation space for salt accumulation in the Zechstein are problematic, as are the extension-related local increases in thickness of Triassic salts. Arguments have been made for varying combinations of rifting, basin flexure and thermal subsidence as the dominant mechanisms controlling salt distribution and thickness (Coward 1995). This is further complicated in the Triassic by namakiers in some parts of the basin floor that were locally recycling dissolved salt into the depositional landscape during episodes of tectonic extension (Mohr et al. 2007). Unlike the transitional to oceanic crust that underlies the Mesozoic postrift Atlantic basins, the continental crust beneath the Zechstein meant that outside of short periods

of Triassic extension or late Cretaceous shortening, basinfloor subsidence across the rest of the 250 Ma history was slow and much less asymmetric in terms of rates of thermally-induced subsidence compared to the longterm history of areas of widespread salt underlain by transitional or ocean crust (Fig. 5.1.12). This prevented rapid asymmetric subsidence into the marine realm and so prevented set up of a depositional system characterised by rapid loading of more landward portions of the salt basin by thick detrital fluvio-deltaic and deep marine depopods. Rather, thinner-bedded salt sedimentation styles of Triassic sequences atop thick halokinetic Zechstein salt indicate that the postZechstein Mesozoic depositional system remained intracontinental to epeiric-marine throughout the Triassic with short episodes of hydrographic isolation and deposition of bedded halites during Buntsandstein, Muschelkalk and early to middle Keuper time (Fig. 5.1.2a). The combination of Triassic basin extension atop a salt décollement, differential loading and locally intense episodes of Zechstein salt flow created Triassic namakiers flowing out onto the land surface (Mohr et al. 2007). In the history of various Zechstein diapirs these salt tongues flowed from the diapir feed-stem not once, but multiple times, to create classic “Christmas Tree” profiles (Fig. 5.1.16a) and outflows of salt into adjacent fault-defined depression (Fig. 5.1.16b). This style of at-surface or nearsurface saltflow created a local highly-evaporitic surface environment and supplied saline brines into the surrounding landscape. Both processes may have aided the deposition of bedded Triassic halite in adjacent playas, Ziegler (1975) argued that much of the Triassic salt may have originated from leached and reprecipitated Zechstein halites, Holser and Wilgus (1981) argued the high bromide content of the halites imply a marine origin, while Fisher

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J.K. Warren Figure 5.1.17. Late Triassic palaeogeography of northwest Europe showing an arid continental interior region dominated by wind-reworked clayey floodplain and emphemeral muddy playa deposits (after Talbot at al. 1994)

and Mudge (1990) suggest that the increase in marine influence eastwards toward the Tethyan marine margin was the ultimate brine source, not halokinetically recycled Zechstein salt. The relative importance of Triassic salt beds can perhaps be better ascertained by comparison with similar modern environments. At the time Triassic salt was deposited much of central Europe was a desert, created at least in part by the highly continental position of the region, which was located well within the interior of the Pangaean supercontinent and in the palaeo-horse latitudes, with Germany located some 20 °N of the Triassic equator (Fig. 5.1.17; Glennie 1986; Ziegler et al. 1993; Edel and Duringer 1997; Sellwood and Valdes 2006). This was the time when the Mercia Mudstone Group was deposited in Great Britain and the Muschelkalk and Keuper were laid down across much of the Central European Basin system (Fig. 5.1.2). At times when the Triassic depositional setting was intracontinental, a lack of laterally consistent bedded thick evaporites in much of Keuper and earlier redbed successions implies water and solute inflow into the various interior areas was insufficient to maintain perennial saline lake in the drainage basin sump. Depositional styles in the Triassic basin interior, away from the marine-influenced desert margins where the bedded evaporites did precipitate, show strong similarities to sediments accumulating in the present day Lake Eyre drainage basin of Central Australia (Fig. 5.1.18). In this situation the basin dries to where much of the sedi-

ment mass is reworked by wind action and little evidence of primary waterlain structures remain in the monotonous silts and loess dominating the redbed association (e.g., Brookfield 2003). At times of relatively higher Triassic sealevel in NW Europe a shallow epeiric seaway covered much of the same basin floor. Such an epeiric system deposited the widespread carbonate mudstones of the Muschelkalk and the only laterally persistent evaporites accumulating at that time were thin CaSO4-nodular beds laid down in sea-marginal mudflats (Borkhataria et al. 2006). Thus, in this Triassic sag basin, times of siliciclastic intracontinental deposition were dominated by a characteristic assemblage of fine-grained facies, the most widespread of which is a variably dolomitic red siliciclastic mudstone, with rare sulphate nodules and little or no halite (Reinhardt and Ricken 2000). Beds are typically massive and unfossiliferous except for occasional vague horizontal laminae and irregularly distributed wispy sand patches (Talbot et al. 1994; Brookfield 2003). Desiccation features are common, as are blocky weathering patterns. Less common are occasional muddy layers with nodules of gypsum/anhydrite and celestine, along with gypsumfilled desiccation cracks. There are also pseudomorphs after displacive halite and suggestions of haloturbation, but no bedded halite is preserved over much of the Triassic interior outside of particular regions. Locally there are also 10-30 cm thick dolomite and limestone beds. These thin beds range from micritic mudstones to ooid grainstones and packstones with freshwater

Salt as sediment in the Central European Basin system as seen from a deep time perspective

algal remains (charophytes and Botryococcus). These local freshwater intervals are interpreted as areas of ponded floodwaters brought to the basin after exceptional rains (Talbot et al. 1994). As for the rest of the Triassic siliciclastic lithotypes, they too have their counterparts in the various siliciclastics sediments of the Quaternary Lake Eyre Basin (e.g., Gibling et al. 1998). Clearly, minor Triassic redbed-hosted evaporites did form in the acid continental groundwater sumps in the interior of this Triassic desert, as in Quaternary redbed playa basins of the Australian interior today, but there was insufficient brine supply to facilitate the deposition and preservation of thick widespread bedded halites in the playa sumps of the supercontinent interior. The same lack of widespread evaporites salts characterises the various playas and salt lakes in the Lake Eyre basin and other continental interior drainage systems in low-relief desert regions of the Quaternary world (see Warren 2006 for summary). For example, in the Quaternary continental interior systems of salt lakes of Australia (29 °S) and the chotts of North Africa (≈33 °N) the lack of a long-term sa-

line brine supply, as well as occasional freshwater floods and the effect of regional upwelling artesian waters, all tend to redissolve and recycle any intracontinental bedded evaporites forming in these drainage sumps (Figs. 5.1.18, 5.1.19; Magee 2004; Bryant et al. 1994). Such intracontinental systems lack a stable brine curtain (Warren 2006). A similar hydrology over much of the Triassic continental interior of Central Europe meant preservation potential for bedded salt was low (Fig. 5.1.17). The floor of Lake Eyre is currently 15 m below sealevel and is the deepest locality on the Australian landsurface (Fig. 5.1.18b). But the Lake Eyre playa and the other ephemeral salt lakes of the southern end of the Great Artesian Basin are separated from a seepage source of marine brine by more than 300 km. The Flinders Ranges and the Gawler Craton create the intervening high. The total thickness of bedded halite in the Lake Eyre sump is less than a few metres. It is completely redissolved every few years by occasional floodwaters sourced in exceptional monsoons rains in Northern Australia. Flood waters then flow more than 1500 km through the Channel Coun-

Figure 5.1.18. Continental interior depositional styles. A) Regional extent of Great Artesian Basin showing flow character with salt lakes in SE part of the basin. B) Landsat image showing deposition in the endoheic Southern Lake Eyre basin is a combination of clayey floodplains, dunes and ephemeral saline lakes. No thick bedded halites are accumulating, only minor sulphate nodules, gypcretes and pseudomorphs (in part after Warren 2006; Landsat image courtesy of NASA)

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J.K. Warren Figure 5.1.19. Chotts of Northern Africa A) Location. B) Regional hydrology showng chotts located at artesian groundwater zone (after Bryant et al. 1994)

try to ultimately gather and evaporate in the southern parts if the drainage basin. The southern area of the Lake Eyre Drainage Basin is also the outflow zone of the Great Artesian Basin of Australia where spring mounds line up along basement faults. This upwelling of artesian water also tends to dissolve evaporites deposited in lakebeds and flushes the resulting solutes back to the lake surface (see Warren 2006; Chapter 3, for more detail and relevant references).

tial edge of a large area of evaporite-lean interior redbed deposits. That is these marine-fed groundwater sumps occupied a geographic position at the hydrologic interface between the Tethyan marine shelf and the evaporitepoor redbeds of the continental interior. Additionally, for a thicker bedded Triassic halite to accumulate, seems to have required an episode of local extension and Zechstein salt withdrawal. This extention formed hydrographically isolated marine-fed groundwater sumps (Fig. 5.1.2a).

The hydrology of Lake Eyre and other low relief arid zone Quaternary deserts show that to preserve more than a few nodules and haloturbated residues requires the presence of subsealevel marine drainage sumps. Such a hydrology can only operate in proximity to a marine mother brine source. This distribution is implied also in the Triassic palaeogeographic reconstruction of the European region (Fig. 5.1.1b) where bedded Triassic halite deposition typifies locally-subsiding depressions near the circumferen-

The only possibility of modern marine waters ever reaching Lake Eyre depression would be if the current inactive rift grabens of Spencer’s Gulf and Gulf St. Vincent were reactivated and propagated northward (Fig. 5.1.18b). These graben-gulfs are the aulocogen remnants of rifts that were active during the Tertiary split of Australia and Antarctica (Veevers 2004). In their current tectonic situation, as part of a passive continental margin at the southern end of the northward drifting Australian con-

Salt as sediment in the Central European Basin system as seen from a deep time perspective Figure 5.1.20. Quaternary namakier/playa analogs at Kuh-e-Namak-Qom WNW of Qom, central Iran (34°45’ 37’’ N, 50° 41’ 18’’ E). A) Dem-draped Aster Image of Kuh-e-Namak-Qom, and the surrounding region (North is toward top of image). The adjacent crescent-shaped playa sits atop the withdrawal sink or rim syncline created by salt flow in the diapir stem. Inset is a 3D representation of part of the same area) B) View from top of the active 300m-high namakier (Miocene mother salt) looking across the playa toward the holy city of Qom. Note that this white saltencrusted playa lies atop the withdrawal sink and is adjacent to the edge of the actively dissolving tongue. C) View from the playa floor toward the side of the active namakier of Kuh-e-Namak Qom showing salt is actively flowing from the vent and out over exposed Miocene sands and muds. D) Dem draped aster image showing cascading namakier tongue and rim of rafted volcanic blocks at its dissolution edge. E) Geology of Kuh-e-NamakQom (interpretation is based on Soder 1952 and field work of author in October 2006)

tinent, this is extremely unlikely. In the Early-Mid Tertiary, when the same subsiding grabens were active, they were located further south in temperate latitudes beneath a humid climate that was unsuited to evaporite deposition. Thus the continental interior of Quaternary Australia is a reasonable analogue for non-evaporite deposition of the continental Triassic interior of the Central European Basin, but not for the local extensional marine-margin episodes when bedded Triassic evaporites were laid down. From a deep time depositional perspective, no Quaternary continental interior basin in any low relief arid region of the world is suitable analogue for deposition of thick Triassic salt beds. But, what if salt can be supplied to a Quaternary hydrology from halokinetic structures that are

actively recycling older salt to in a desert landscape. Such systems are active in the interior deserts of Iran where varying combinations of Miocene or Infracambrian salt now outcrops as active salt glaciers. Ziegler (1975) argued a similar system supplied solutes for Triassic salts from a halokinetically recycled Zechstein source. In the Qom Basin west-northwest of the holy city of Qom, Miocene salt of the Lower Red Formation is actively cascading out onto the surface as a namakier or salt glacier (Fig. 5.1.20a). This salt was first laid down as marinefed salt in extensional basins of the Oligo-Miocene Lower Red Formation (Rahimpour-Bonab et al. 2007). The salt fountain at Qom Kuh rises to 1235 m asl and rises 315 m above the surrounding plateau (Fig. 5.1.20a). It forms

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J.K. Warren Figure 5.1.21. Degrading namakiers sourced in Oligo-Miocene salt, Central Iran. A) Kuh-e-Namak-Lash (34° 20’ N, 51° 15’ E) shows retreating salt apron covered by foundered blocks of Qom Fm and a peripheral apron of Eocene volcanics, some of which now lie largely outside the salt apron. B) Kuh-e-Gach-Saveh (34° 53’ 53’’ N, 50° 11’ 31’’). The position of the formerly active salt vent is indicated by a gypsum carapace atop a layer of insoluble residues. The former position of the salt tongue is define by a zone of transpression at the edges of a lens-shaped fragmented apron made up of disturbed blocks of Qom Fm limestone. Facies patterns in bedded Qom Fm and underlying Upper Red Fm some 1.5 km east of vent suggest this tranpressional structure is centred on an earlier Oligo-Miocene extensional salt structure (interpretation is based on author’s field work). Kuh-e-namak means mountain of salt; Kuh-e-Gach means mountain of gypsum. A geographic term should always be appended to the name to avoid confusion as there are many examples of both outcropping in Iran

a smooth hemispheric dome with a diameter of 2.5 km, capped by a 0 -3 m thick layer of pocked salt dissolution residues and clays (Fig. 5.1.20b). The namakier tongue is surrounded and underlain by a dissected apron of Upper Red Formation and Recent sediments. Figure 5.1.20c is a view of the eroded side of the extruded salt tongue showing it overflows a tilted collar of Upper Red Formation sediments. The active but buried diapiric vent is located beneath smooth dome mound defining highest part of the salt tongue (Fig. 5.1.20d). This cascading geometry indicates salt is actively spewing from this vent. A similar nippled salt mound at Kuh-e-Namak-Fars shows that this vent is also active, but in this case the age of the flowing salt is Infracambrian, not Oligo-Miocene (Fig. 5.1.22b) Mapping of the active salt flow at Kuh-e-Namak Qom (Fig. 5.1.20e) and nearby, now largely inactive at-surface salt structures at Kuh-e-Namak-Lash and Kuh-eGach-Saveh (Fig. 5.1.21) shows that the at-surface salt

in all three structures is sourced in the Oligo-Miocene Lower Red Formation. The equivalence of halokinetic stratigraphy between salt structures in the eastern Kavir, where the halite comes from a combination of Lower Red and Upper Red Formations, and salt structures in the Qom area, as inferred by Talbot and Aftabi (2004), does not hold true in terms of regional outcrop, dissolution breccia and subsurface geology. Flowing salt in all three at-surface salt structures in the Saveh-Qom-Lash area comes from halite in the Lower Red Formation where it both lubricates regional thrusts and is caught up in now-squeezed diapirs that were first active in an extensional phase that also influenced the facies distributions in deepwater bryozoal carbonates of the overlying Qom Fm. Flowing salt related to subsequent Plio-Pleistocene compression has carried large blocks of Eocene volcanics to the surface at both Kuh-e-Namak-Qom and Kuh-e-Namak-Lash. Gravity sliding carries these blocks of denser material to the periphery of the salt mass where

Salt as sediment in the Central European Basin system as seen from a deep time perspective

they now form a debris apron about the dissolving perimeter of the salt tongues. The three salt structures (Qom, Lash and Gach) illustrate what happens to an at-surface namakier in a desert setting as salt flow into the at-surface structure slows and ultimately ceases. The smooth fountain profile at Kuh-eNamak-Qom shows that salt is still issuing from this structure (Fig. 5.1.20d). In contrast the surface profile across the salt tongue at Kuh-e-Namak-Lash is very irregular as the structure is now capped by large floundered blocks of Qom Fm overburden (Fig. 5.1.21a). Salt flow has largely ceased and the tongue edge is retreating. At Kuh-e-Gach-Saveh, the most evolved structure, at surface halite is nowhere to be seen and a gypsum carapace sits atop a zone of widespread insoluble residues that mark the position of a former halite tongue and vent (Fig. 5.1.21b). The lack of at-surface halite and the adjacent down-dip terrain of fragmented Qom Fm limestone blocks shows that this namakier is in at even more mature stage of its weathering than Kuh-e-Namak-Lash. This time/flow related evolution, from active namakier to a layer of clay residuals, illustrates low preservation potential for any desert namakier once the salt supply/renewal ceases. In all three cases ongoing dissolution of at-surface salt is contributing solutes to the surrounding landscape. A

salt-encrusted saline playa plain has formed atop a small depression created by the active episode of salt flow that is supplying Qom-Kuh. In the case of the other two tongues (Lash-Kuh and Gach-Kuh) the dissolved solutes are dispersed to the landscape and no playa is located nearby. It seems the climate over the Qom-Kuh region, even as salt is flowing is not sufficiently arid, nor the volume of redissolved halokinetic salt sufficient to form thick beds of halite in any nearby groundwater sumps or depressions. If we look at other active or recently active namakiers in Iran sourced in Oligo-Miocene salt in the Eastern Kavir such as Dome 10 as studied by Jackson et al. 1990 or the various Kuh-e-Namaks in Fars Province that are sourced in Infracambrian Hormuz salt (summarised in Chapters 6 and 7 of Warren 2006) we see a similar style of salt tongue retreat. Even in the largest active Hormuz-salt sourced namakiers no more than thin ephemeral halite crusts characterise the braid plains and playas in adjacent withdrawal sinks (Fig. 5.1.22b). Although namakiers in today’s deserts are best developed in the arid regions of Iran, all show signs of rapid retreat and dissolution once the supply of salt is cut off. None seem to have longterm preservation potential. But from a deep time perspective we know that namakiers sourced in Permian Zechstein salt were at the landsurface and subaerial a number of times in Triassic history of the Figure 5.1.22. Namakier (28°N, 54° 55’ E) sourced in Infracambrian Hormuz salt in the Fars region of the Zagros fold and thrust belt in SE Iran (Landsat image courtesy of NASA)

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Central European Basin system and that they had sufficient volumes to be preserved. Without a deep time perspective strict actualistic uniformitarianism would argue that salt tongues situated at the surface in a continental interior basin have no preservation potential. The presence of salt tongues flowing across the Triassic landscape of parts of NW Europe is a direct reflection of the lack of equivalence between the relatively milder conditions of today’s continental interior desert climates and the much more arid continental interior climates of the Pangaean supercontinent (Mohr et al. 2007). But even

in the hyperarid conditions of the Triassic it probably also required a marine margin seepage setting to accumulate thick widespread-bedded halites. Thick widespread evaporite beds, away from any local namakier source, are not directly related to meteoric water inflow or to a recycling of older salt into a continental landscape, but are tied to the set up of an appropriate marine margin subsealevel seepage hydrology. The present icehouse climate mode of the earth and a lack of a supercontinent-induced interior hyperarid climate both serve to make the present a very limited key for explaining the scale and diversity of past evaporites.

5

Chapter 5.2

Flow and Transport Properties of Salt Rocks J.L. Urai · Z. Schléder · C.J.Spiers · P.A. Kukla

5.2.1 Introduction Ductile evaporites play a key role in controlling the dynamical evolution of many sedimentary basins. We review the mechanical and transport properties of these rocks focusing on halite, bischofite and carnallite. Reliable modelling of salt flow during basin evolution, or of salt flow related to long-term engineering challenges, requires extrapolation of experimentally-derived flow laws to strain rates much lower than those attainable in the laboratory. This extrapolation must be based on an understanding of the microscale deformation mechanisms operating under these conditions, obtained by combining studies of natural laboratories with experimental work. The engineering creep laws generally

used in the salt mining industry are based on dislocation creep processes quantified in laboratory experiments of necessarily limited duration. However, a large body of evidence clearly demonstrates that under conditions of long-term deformation, grain boundary dissolution-precipitation processes, such as solution-precipitation creep (or “pressure solution”) and dynamic recrystallisation, play a significant role. The operation of these processes can cause major changes in rheology. Moreover, the high fluid pressures associated with deforming evaporite systems can lead to dramatic increases in permeability, strongly reducing sealing capacity. These properties must be incorporated in quantitative models of evaporite basins to obtain realistic descriptions of salt behavior at the necessary range of length and time scales.

Table 5.2.1 List of the main evaporite minerals and the wireline log properties of evaporite rocks formed by these.

Name

Formula

Sonic transit time

Density

GR

Neutron “Porosity”

kgm-3

API

%

msft-1

Bischofite

MgCl2 . 6 H2O

1560

0

> 60

100

Carnallite

KMgCl3 . 6 H2O

1570

220

65

78

Epsomite

MgSO4 . 7 H2O

1710

0

> 60

Sylvite

KCl

1860

500

-3

74

Halite

NaCl

2040

0

-3

67

Kainite

MgSO4KCl . 3 H2O

2120

245

45

Gypsum

CaSO4 . 2 H2O

2350

0

>60

Kieserite

MgSO4 · (H2O)

2590

0

38

Calcite

CaCO3

2710

0

-1

49

Polyhalite

K2Ca2Mg(SO4)4·2 (H2O)

2790

180

15

57

Langbeinite

K2Mg2(SO4)3

2820

275

0

52

Dolomite

CaCO3 MgCO3

2870

0

1

44

Anhydrite

CaSO4

2980

0

-2

50

52

278

J.L. Urai · Z. Schléder · C.J. Spiers · P.A. Kukla Figure 5.2.1. Schematic drawing of the microstructural processes that can operate during deformation of rock salt at temperatures in the range 20200 °C. Different shades of green represent crystals with different orientation. See text for explanation

5.2.2 Physical properties of evaporites Most evaporite rock sequences consist mainly of carbonates, sulphates and chlorides, Halite rock is generally the dominant chloride lithology and due to its low creep strength, low static porosity, permeability and low density it exerts a profound influence on basin evolution - through salt tectonic movement and fluid trapping. Other chloride evaporite rocks, though volumetrically much less important than halite, carbonate or sulphate lithologies, can also have important effects because of their exceptionally weak rheology (causing heterogeneous deformation and drilling problems) or chemical composition (salt metamorphism and fluid flow). In table 5.2.1 the main evaporite minerals are listed together with their wireline log properties.

5.2.3 Deformation mechanisms and rheology of halite in experiments 5.2.3.1 Deformation mechanisms and associated processes Polycrystalline halite rocks (rock salt) consist of grains of halite (NaCl), with a diameter between 0.01 mm and

several dm, containing impurities in solid solution, secondary mineral phases and fluids trapped in inclusions, grain boundaries or in pores. The mean grain size of most halite rocks lies in the range 2-20 mm, though extrusive salts may show much finer mean values (0.5 mm) while secondary salt can be much coarser (10-30 cm). Under deviatoric stress, rock salt can deform by a range of processes. The deformation mechanisms known to operate at temperatures relevant for engineering and natural halokinetic conditions (20-200 °C) are summarised in figure 5.2.1. At very low effective confining pressures (less than a few MPa) and high deviatoric stresses (> 1520 MPa), inter- and intragranular microcracking, grain rotation and intergranular slip are important strain accumulating processes alongside crystal plasticity, and the mechanical properties and dilatational behaviour are dependent on the effective mean stress or effective confining pressure (Cristescu and Hunsche 1998; Peach and Spiers 1996; Cristescu 1998; Peach et al. 2001). At sufficiently high deviatoric stress, the material fails in a (semi)brittle manner, with failure described by a pressure (effective mean stress) dependent failure envelope. With increasing effective mean stress, microcracking and dilatancy are suppressed and crystal plastic processes dominate. At temperatures in the range 100-200 °C, dislocation creep is important in laboratory experiments, and polycrystalline halite can deform to large strains by this mech-

Flow and Transport Properties of Salt Rocks Figure 5.2.2. Reflected light optical micrograph of experimentally deformed rock salt (Asse Speisesalz, 150 ºC, 50 MPa, 3x10-5 s-1), showing dislocation slip lines, incipient subgrains and minor grain boundary microcracking. Image is approximately 1.5 mm wide

anism (Fig. 5.2.2), even at confining pressures as low as 10 MPa. During this process, subgrains are formed in the halite grains (Pennock et al. 2005, 2006a,b), with the diameter of the subgrains showing an near-linear, inverse correlation with deviatoric stress (Carter et al. 1993, see Fig. 5.2.11). If the polycrystal contains small but significant amounts of water in the form of saturated brine inclusions or grain boundary films, as is generally the case for both natural and synthetic samples, fluid assisted grain boundary migration generally operates. This is an efficient process of reducing dislocation density and hence removing the stored energy of dislocations, even at room temperature (Schenk and Urai 2004; Schenk et al. 2006, see Fig. 5.2.3). While dislocation creep processes take place in the crystal lattice of the halite grains, and fluid assisted grain boundary migration involves solution-precipitation trans-

fer across grain boundaries, solution-precipitation creep, or “pressure solution”, is a process which involves mass transfer around grain boundaries. Here, in the presence of a small amount of saturated grain boundary brine, grains dissolve at highly stressed boundaries, and after diffusion of the material through the grain boundary fluid, the material crystallises at interfaces under low normal stress (Schutjens and Spiers 1999; Spiers et al. 2004; Fig. 5.2.4). This process is accompanied by intergranular sliding and rotation (grain rearrangement), and can lead to compaction of porous salt or to deviatoric strain of non-porous aggregates (Spiers et al. 1990). Solution - precipitation creep is an important deformation mechanism in most wet rock systems in the Earth’s crust (Renard and Dysthe 2003), but is especially rapid in rock salt. Early reports, theoretical treatments and

Figure 5.2.3. Reflected light micrograph of experimentally deformed rock salt, showing deformed grains replaced by new, strain free grains (Asse Speisesalz, 150 ºC, 100 MPa, 3x10-5 s-1 followed by stress relaxation). The grain boundary migration is assisted by the presence of thin fluid films on the grain boundaries, and can take place at significant rates at room temperature. Image is approximately 1.5 mm wide. b) Diagram illustrates the principle of grain boundary migration by solution-precipitation transfer across fluid-filled grain boundaries

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Box 5.2.1 The recrystallisation process involves grain boundary migration by solution-precipitation transfer across grain boundary water/ brine films (Fig. 5.2.3b), and is driven by chemical potential differences across grain boundaries related to the dislocation density differences Δρ between old deformed grains and newly growing grains (Peach et al. 2001). In strongly deformed, wet rock salt, the migration process is very rapid, reaching rates up to 10 nms-1 at room temperature (Schenk et al. 2006). A model for the process has been derived by Peach et al. (2001), and predicts the following equations giving grain boundary migration velocity V, depending on whether cross-boundary diffusion or the interfacial reactions of dissolution and precipitation are rate controlling: Diffusion control

(5.2.1)

Interface control

(5.2.2)

Here, Dgb (m2s-1) is the diffusivity of the solute (NaCl) in the grain boundary fluid, C is its solubility (m3m-3), t is the fluid (brine) film thickness (m), W is the self-energy per unit length of stored dislocations, Ω is the molar volume of the solid phase (NaCl), R is the gas constant, T is absolute temperature and I is the linearised rate coefficient for dissolution and precipitation from the fluid film. The diffusion controlled model predicts rates of grain boundary migration that agree favourably with observations on rock salt (Peach et al. 2001).

Figure 5.2.4. (a) Typical microstructure of solution-precipitation deformation in a porous halite polycrystal containing saturated brine (after Spiers et al. 1990), t marks sites of truncation of grains by pressure solution. Diagram shows the mechanisms of solution-precipitation creep without plasticity of the crystals (b) via a thin fluid- filled grain boundary, and (c) combined operation of solution-precipitation creep and crystal plasticity

reviews of solution-precipitation creep in rocks are given by Durney (1976), Rutter (1976), Sprunt and Nur (1977), Rutter (1983) and Tada and Siever (1996). Recent theoretical treatments of the process are given by Lehner (1990) and Kruzhanov and Stöckhert (1998). In brief, the differences in chemical potential μ between points in the

solid at grain boundaries under high stress and those under lower stress provide the driving force for dissolution, transport by diffusion in the intergranular fluid, and precipitation (Fig. 5.2.3). Additional driving force (chemical potential drop) both along and across grain boundaries can be provided by internal plastic deformation of the

Flow and Transport Properties of Salt Rocks

grains, giving rise to combined grain boundary migration and solution-precipitation creep. The above processes have been documented in laboratory experiments and in naturally deformed salt from a wide range of settings (Urai et al. 1987; Trimby et al. 2000; Ter Heege et al. 2005a,b; Schléder and Urai 2007; Schléder et al. 2007). The relative importance of each process depends strongly on variables such as temperature, confining pressure, grain size, solid solution impurities and second phase content, and, importantly, on the presence of sufficient water in grain boundaries to enable solutionprecipitation phenomena (Fig. 5.2.5). Fluid assisted grain boundary migration and solution-precipitation processes do not operate in dry salt, i.e., synthetic samples made of carefully dried material (< 5 ppm water, Ter Heege et al. 2005a,b) or natural samples deformed under dilatant conditions which allow intergranular water to evaporate (Peach et al. 2001). The fields in which different deformation mechanisms are dominant can be represented in deformation mechanism maps (Ter Heege et al. 2005a,b). Note that published deformation mechanism maps tend to differ in the mechanisms included, depending of the time scales of interest and on whether the effects of water (brine) are included or not.

rate of interest. For rock salt, summaries of behaviour observed in experiments have been published by, for example, Urai et al. (1986b), Cristescu and Hunsche (1998), Hunsche and Hampel (1999) and Ter Heege et al. (2005a,b).

5.2.3.2 Rheological behaviour – “flow laws”

Solution- precipitation creep has been widely recorded in laboratory experiments on wet, fine grained (< 500 μm) polycrystalline halite at temperatures in the range 20200 °C, leading to quite rapid linear viscous deformation of dense salt (Eq. 5.2.2) and compaction of porous material (e.g., Urai et al. 1986; Spiers et al. 1990; Renard et al. 2004; Ter Heege et al. 2005b). The process has also been observed at individual halite grain contacts under stress (e.g., Gratier 1993; Schutjens and Spiers 1999; De Meer et al. 2005) and in salt aggregates containing second phases (Renard et al. 2001; Zoubtsov et al. 2004). Possible evidence for pressure solution has also been reported by Berest et al. (2005) in low stress creep experiments on coarse, natural rock salt, in which much faster rates of deformation were observed than expected by extrapolating conventional dislocation creep laws obtained at higher stresses and strain rates (cf. Fig. 5.2.5). In addition, rapid long-term deformation of pillars and galleries in potash mines has been attributed to solution-precipitation creep in rock salt (Campos de Orellana 1998; Lee and de Souza 1998), though such in-situ experiments have not yet been sufficiently documented to allow full quantification of the processes operating (Bekendam and Urai 2007).

The rheology of a given crystalline material depends on the dominant deformation mechanism, which in turn depends on the time scale and hence deformation

However, because of the strong grain size dependence and limited duration of laboratory tests, solution-precipitation creep is rarely seen in experiments on natural rock Figure 5.2.5. Differential stress - strain rate diagram summarising low temperature laboratory data for a wide range of salts, from experiments by Sandia, BGR, Utrecht University and other Laboratories. Application of these data to conditions of long-term creep requires extrapolation. Broken line is extrapolation of the (grain size insensitive) dislocation creep law, taking n = 5 (BGR creep law, 30 - 50 ºC). Solid line is the room temperature solution-precipitation creep law for different grain sizes. It can be seen that for a grain size of 3 mm the two mechanisms contribute equally to the deformation at a strain rate of 10-10 to 10-11 s-1. Time to reach 10% strain at different strain rates is shown on the right side of the graph (after Schléder and Urai 2005)

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Box 5.2.2 Considering steady state, non-dilatant deformation, the main classes of equations for the creep strain rate of rock salt are written for dislocation creep, and

(5.2.3)

(5.2.4) for solution- precipitation creep, with the total strain rate being the sum of the two (5.2.5) In these equations, written in a form appropriate for axially symmetric compressive deformation, A and B are material parameters, QDC and QPS represent (apparent) activation energies for dislocation and pressure solution creep, R is the gas constant, T is absolute temperature, σ1 andσ3 are the maximum and minimum principle compressive stresses, D is grain size, and n and m are the exponents of stress and grain size respectively. Two important differences between equations 5.2.1 and 5.2.2 are firstly the dependence of strain rate on stress (n = 1 for solution-precipitation creep or pressure solution while n > 1 for dislocation creep – see figure 5.2.5 and 5.2.6), and secondly the dependence of strain rate on grain size. Note that for dislocation creep deformation is grain size independent, while the exponent m = 3 makes pressure solution creep strongly grain size dependent (Fig. 5.2.5 and 5.2.7). At 20-200 °C, differential stresses below 15-20 MPa and strain rates below 10-6 s-1, both uniaxial and triaxial experiments on natural and synthetic rock salt show power law dislocation creep behaviour (Eq. 5.2.3) with a stress exponent n of 5-6 at the higher stresses and 3.5-4.5 at lower stresses (Wawersik and Zeuch 1986; Carter et al. 1993; Hunsche and Hampel 1999). The apparent activation energy for creep is unusually low, taking values of 50-80 kJmol-1. Intragranular microstructural signatures including wavy deformation band (slip/sub-boundary) structures suggest that cross-slip of screw dislocations may be the rate controlling process at differential stresses (σ1-σ3) above 10-15 MPa (n = 5-6), while well formed equiaxed subgrains indicate that climbcontrolled recovery becomes dominant at lower stresses (n = 3-4)solid solution, amount and distribution of secondary mineral phases, grainsize, subgrain size, dislocation density and fluids in grain boundaries. Deformation experiments, performed in the dislocation creep field at confining pressures high enough to suppress dilatancy (> 10-20 MPa), have shown that “wet” samples containing more than 10-20 ppm of water (brine) at grain boundaries undergo rapid dynamic recrystallisation by fluid assisted grain boundary migration, alongside dislocation creep. Compared with dry rocksalt samples (<10 ppm water), this recrystallisation process leads to a reduction in average dislocation density, an associated reduction of flow strength by 25-50% and power law flow behaviour with an n-value of about 4.5 (Peach et al. 2001; Ter Heege et al. 2005b). In addition, during dynamic recrystallisation of wet salt (>10-20 ppm), the grain size evolves such that a systematic relation between grain size, flow stress and temperature is established, with deformation occurring close to the boundary between the dislocation and solution-precipitation or pressure solution creep fields (Ter Heege et al. 2005a; Fig. 5.2.8). The equation relating mean recrystallised grain size (DDRX measured in μm) to differential stress (MPa) and temperature obtained by Ter Heege et al. (2005a) shows only weak temperature dependence and can be viewed as a basis for estimating flow stresses from the dynamically recrystallized grain size of natural salt (i.e., as a so-called palaeopiezometer). The relation is given by (5.2.6) where b is the mean Burgers vector for halite (b=3.99 x 10-4 μm), G is its shear modulus (G=1.5 x 104 MPa), log(K) = -1.55±0.24, p = 1.85 ± 0.23 and the (apparent) activation energy term QDRX/a = 14.2±2.8 kJmol-1.

salt (grain size typically 1 cm), and it is therefore not usually included in engineering descriptions of salt rheology (Cristescu and Hunsche 1998; Hunsche and Hampel 1999). Nonetheless, predictions made using equation 5.2.4 suggest that provided the salt contains sufficient water (>10-20 ppm, as most natural salts do), pressure solution creep should become important at strain rates below those reached in experiments (see Fig. 5.2.5 and 5.2.6). As indicated above, at low confining pressures and high deviatoric stresses, flow of rocksalt is accompanied by

dilatant grain boundary microcracking and rapid permeability increase (Cristescu and Hunsche 1998; Peach and Spiers 1996). The mechanical conditions under which this occurs have been accurately delineated by Cristescu and Hunsche (1998) and Cristescu (1998). While the onset of microcracking has a minor direct effect on creep behaviour, it is important to note that it can strongly influence the effects of water on creep. In salt containing small quantities of water, microcracking disrupts grain boundary films and inhibits both grain boundary migration and pressure solution, particularly if the water can

Flow and Transport Properties of Salt Rocks Figure 5.2.6. Deformation mechanism map for dense rock salt, incorporating solution- precipitation creep for a grain size D = 10 mm. It can be seen that at low temperature and slow strain rates deformation is in the transition between dislocation creep and solution- precipitation creep (after Spiers et al. 1990)

escape from the sample (Peach et al. 2001). On the other hand, under conditions where microcracking allows free brine or water vapor access to the interior of a creeping salt sample, then both recrystallisation and solution-precipitation creep effects can be strongly enhanced. Note that despite the large amount of data now available on solution-precipitation creep in salt, details of the

microphysics of the process are incompletely understood. This is at least partly due to the difficulties of imaging the fine-scale (1-200 nm) structure of wetted grain boundaries during deformation. Approaches applied here include in-situ optical, infrared and electrical resistivity measurements, interference microscopy and electron microscopy of frozen boundaries using cryo-SEM (Hickman and Evans 1995; Schutjens and Spiers 1999; Spiers et al. Figure 5.2.7. Dynamically recrystallised grain size versus stress data for synthetic rock salt samples, superposed on a deformation mechanism map, showing that these samples deform in the transition region between dislocation creep and solution- precipitation creep (after Ter Heege et al. 2005a,b)

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J.L. Urai · Z. Schléder · C.J. Spiers · P.A. Kukla Figure 5.2.8. Steady state strain rate data of finegrained wet rock salt samples deforming by solution- precipitation processes, showing a stress exponent n close to 1. Gray bands represent the theoretical flow law for this process (Eq. 5.2.4). (after Spiers et al. 1990)

2004; De Meer et al. 2002, 2005; Schenk et al. 2006). Insitu methods applied during active pressure solution in salt indicate that the grain boundary fluid is contained in a non-equilibrium or dynamically wetted island-channel structure, whose thickness varies with the crystallographic orientation and charge of the bounding crystal surfaces (e.g., De Meer et al. 2005). At sufficiently low intergranular effective stresses, however, surface energy driving forces are expected to become large enough, compared to stressrelated driving forces, to cause healing of grain boundaries by island or contact growth, forming isolated fluid inclusions in the boundary (Schutjens and Spiers 1999). A basis to evaluate the conditions under which this will occur has recently been investigated by Van Noort et al. (2007) but is as yet difficult to apply in practice to a material such as halite. Further work is still needed in this area.

5.2.4 Deformation mechanisms and rheology of carnallite and bischofite Much less is known about the rheology and deformation mechanisms of other evaporites. Carnallite and bischofite, which can both form layers up to 30 m thick, were inves-

tigated by Urai (1983), van Eekelen et al. (1984), Urai and Boland (1985), Urai (1985), Urai et al. (1986a), Urai (1987a,b) and Schenk and Urai (2005). It was shown that there are many similarities with halite, and dislocation creep, dynamic recrystallisation and solution-precipitation processes (Fig. 5.2.9) have been shown to occur both in laboratory samples and in naturally deformed samples. In addition, deformation twinning was shown to occur in both materials. Mechanical data (Fig. 5.2.10) point to a much weaker rheology than halite, while bischofite is much weaker than carnallite. The rheology of anhydrite under conditions of natural deformation is virtually unknown. Existing studies are all done at high temperature and in dry samples, e.g., Dell‘Angelo and Olgaard (1995), and it is not clear how these can be extrapolated to natural conditions.

5.2.5 Natural laboratories Studies of rock salt deformation in nature are essential for reliable extrapolation of laboratory data to describe the flow of salt during slow, human-induced or natural flow, because such studies provide a detailed understanding of the deformation mechanisms and microstructural processes

Flow and Transport Properties of Salt Rocks Figure 5.2.9. Mechanical twinning, subgrain rotation and grain boundary migration in experimentally deformed Asse carnallitite. Width of image is 4 mm

that operate at strain rates well below those accessible in laboratory experiments. In recent years, major advances in this field have been reported, based on developments in microstructural and textural/orientation analysis using electron backscatter diffraction (EBSD), microstructure decoration by gamma-irradiation, Cryo-SEM and other methods. Samples from a wide range subsurface and surface locations have been studied (e.g., Schléder and Urai 2005, 2007; Schléder et al. 2007).

ly, as a function of grain size, impurity content, stress path and fluid chemistry. Differential stress, as measured using laboratory-calibrated subgrain-size piezometry (Fig. 5.2.11), is usually less than 2 MPa in rock salt deforming in nature, in agreement with in-situ stress measurements and geologic flow rates (Schléder and Urai 2005). Higher stresses, up to 5 MPa are recorded in the near-surface parts of diapir stems where salt is extruded to the surface (Schléder and Urai 2007).

In addition there has recently been much progress in measuring the surface displacement field in areas of active salt tectonics, in salt mining districts, on sediment rafts above mobile salt, on emerging salt diapirs, and in areas where removal of ice sheets has led to a change of overburden load. These data can be inverted using nonlinear finite element techniques, to obtain constitutive equations for salt flow during slow natural deformation (Weinberger et al. 2006; unpublished data). Insight into the in-situ rheology of salt on the time scale of years has also been gained by simply adjusting the flow laws used in numerical models of mining-related subsidence to obtain a match with surface displacement evolution. Studies of this type, conducted in relation to deep solution mining operations at Barradeel in the Netherlands (2-3 km depth), suggest salt flow behaviour involving dislocation creep (n value of 3-4) combined with a linear viscous flow law comparable to that expected for solution-precipitation creep.

Microstructural studies of subgrains and recrystallised grains in naturally deformed rocksalt also show, in agreement with recent experiments, that during fluid-assisted dynamic recrystallisation of salt in nature (water content >10 ppm), the grain size adjusts itself so that the material deforms close to the boundary between the dislocation and pressure solution creep fields. Power law flow, as measured in recrystallising samples (Ter Heege et al. 2005a,b), with an n-value of about 4.5 is therefore proposed to be a good representation of this behavior. In samples which are sufficiently fine grained, solution-precipitation creep (Eq. 5.2.2), is found to be dominant both in salt glaciers recrystallised after extrusion to the surface, and in very fine grained primary rock salt in the subsurface (Schléder and Urai 2005, 2007). At geologic strain rates, such salt will be orders of magnitude weaker than would be predicted from extrapolation of short-term experiments on coarse-grained rock salt (see Fig. 5.2.5).

Microstructural studies of naturally deformed salt show that low temperature dislocation glide and dislocation creep processes, solution-precipitation creep and waterassisted dynamic recrystallisation are all of major importance (Urai et al. 1987; Schléder and Urai 2005, 2007). The relative importance of these processes varies strong-

The rather high variability of flow strength in layers of rocksalt in nature is in good agreement with the smallscale folding ubiquitously observed in layers of naturally deformed salt. This has not yet been incorporated in numerical models of salt tectonics, which typically assume much more homogeneous material properties and accordingly produce much less heterogeneous strain fields.

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fluid films to form isolated fluid inclusions (or possibly some other structural change), thus rendering the boundaries immobile in a manner analogous to that proposed above for cessation of pressure solution at low stresses. In the following, we consider a number of recent examples of how microstructural studies of natural salt can elucidate operative deformation processes and rheology in nature.

5.2.5.1 Evidence for diffuse dilatancy and fluid flow in rock salt in the deep subsurface

Figure 5.2.10. Summary diagram of the mechanical properties of bischofite and carnallite, compared with rock salt. It can be seen that carnallite is much weaker than halite, and bischofite in turn is much weaker than carnallite

It is an interesting and as yet unexplained microstructural observation that despite the high rate of fluid-assisted grain boundary migration observed in experiments, most naturally deformed rock salt is not completely recrystallised and preserves subgrains. A possible explanation for this is that below some critical difference in driving force for crossboundary solution-precipitation transfer, surface energy driving forces cause necking or healing of grain boundary

Intact rock salt has an extremely low permeability (< 10-21 m2), which imparts excellent sealing capacity. However, we know from experiments that at sufficiently low effective mean stress, dilatancy and extensional (Mode I) fracture can occur in deforming salt, producing significant permeability (Peach and Spiers 1996). If this occurs in nature, fluid flow will be possible. Evidence for fluid flow, in the form of saturated brines of different chemical composition, and of hydrocarbons is frequently found in naturally deformed salt (Schléder et al. 2007; Schoenherr et al. 2007b; Magri et al. this volume). Microstructural evidence suggests that both diffuse dilatancy and Mode I fracturing can occur at very low effective stress conditions, i.e., at near-lithostatic fluid pressures and differential stress of a few MPa, in the deep subsurface, followed by crack healing through precipitation form supersaturated solutions.

Figure 5.2.11. Subgrain-size versus differential stress data from experimentally deformed rock salt, providing the basis for measurement of in-situ differential stress in core samples. Solid dots with error bars are the application of this technique to Hengelo rock salt (after Schléder and Urai 2005)

Flow and Transport Properties of Salt Rocks

5.2.5.2 Fluid Flow in Fractures: A case study of the Neuhof Mine Germany Zechstein (Z1) rocksalt core samples from the immediate vicinity of the Hessen potash bed from the Fulda basin were studied by Schléder et al. (2007). Here the anhydrite and halite are folded into tight, isoclinal folds and (the sequence) is cut by an undeformed, 1 cm thick, coarsegrained halite vein (Fig. 5.2.12). Microstructures were investigated in etched, gamma-irradiated thin sections from both the halite wall rock and from the vein. The lack of dissolution structures and the widespread occurrence of plate-shaped and hopper grains in the folded halite wall rock suggests that the sedimentary environment was a perennial lake. Deformation microstructures in this folded halite wall rock are in good agreement with the solution-precipitation creep process (5.2.13). In-situ palaeo-differential stress values are not available for the samples because of the absence of subgrains. This implies however that the differential stress was below 0.3 MPa, because otherwise the fine-grained wet halite described (grain size D = 0.5-1 mm) would have developed subgrains. Model calculations suggest that this fine-grained salt is very weak as compared to domal salt and it deforms relatively fast ( ε~ 5x10-10 s-1 for T = 353 K and D = 0.5 mm) even at low differential stresses (σ1-σ3 = 0.1-0.3 MPa). Strength variations in anhydrite-rich and anhydrite-poor layers are accounted for the strong folding in the halite beds. The vein is completely sealed and composed mainly of euhedral to subhedral halite grains, which often overgrow the wall rock grains. Those microstructures, together with the presence of occasional fluid inclusion bands suggest that the crystals grew into a solution-filled open space.

Based on considerations on the maximum value of in-situ differential stress and dilatancy criteria discussed above, and on the amount of released fluids from the potash bed during metamorphism and the volume change, it is proposed that the crack was generated by hydrofracturing of the rocksalt due to the presence of the salt metamorphic fluid at near-lithostatic pressure.

5.2.5.3 Deformation mechanisms in weakly deformed bedded salt in Hengelo, the Netherlands As an example of a study of deformation mechanisms in natural salt, we briefly review the work done recently by Schléder and Urai (2005) on the bedded salt mined by AKZO at Hengelo in the Netherlands. Deformation of the Hengelo rock salt in the geologic past has taken place probably during Cretaceous tectonic inversion in the area. The microstructure of core samples from the subhorizontal, bedded Main Röt Evaporite Member (AKZO well 382, depth interval of 420-460 m) was studied by transmitted and reflected light microscopy of gamma-irradiation decorated samples. Primary microstructures compare favourably with those found in recent ephemeral salt-pans. In addition, in all layers the grains are rich in deformationrelated substructures such as slip bands and subgrains indicating strains of a few percent (Fig. 5.2.14). The study of gamma-irradiation decorated thin sections showed that the main recrystallisation mechanism was grain boundary migration (Fig. 5.2.14). This process removes primary fluid inclusions and produces clear, strain-free (subgrain- and slip band free) new grains. Differential stresses as determined by subgrain size piezometry were 0.45 – 0.97 MPa (Fig. 5.2.11).

Figure 5.2.12. Scanned image of the studied core interval from the Neuhof mine, together with the hand-drawing of traced anhydrite layers (black lines). The transparent layer, which cut through the folded layers, is a coarse-grained, halite-filled vein. Note that there is a slight offset in the wall-rock across the vein. After Schléder et al. 2007

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Figure 5.2.13. A) gamma-irradiated thin section from Neuhof scanned with transmitted light. The E-W trending vein is filled with large euhedral grains. Anhydrite and potash mineral is also intercalated (occur as white patches). The places of detail images of B, C and D are indicated with rectangles. B) detail of the wall-vein interface. Note that the crystals are overgrowth of the wall rock grains. The N-S oriented cracks are due to drilling and sample preparation. The white material patch indicates the former place of the potash, which has been dissolved during sample preparation. C) detail of the contact of the coarse grained vein halite and the wall rock. The elongated white lines are grown-in subgrains. Note the growth bands which are parallel to the {100} crystal facets. At the right side, the white patch shows the former place of the intercalated potash. D) additional image illustrating the vein-wall interface. Note that the vein crystals overgrow the wall-rock grains

The corresponding deformation mechanisms are inferred to be a combination of dislocation creep, dynamic recrystallisation and solution- precipitation processes. Solution-precipitation processes are activated by the small amount of brine present in grain boundaries. Inserting the above differential stress values determined from subgrain size into published flow laws for dislocation creep and for pressure solution (Carter et al. 1993) yields transitional strain rates between 10-12 and 10-13 s-1. The data provide a view of very slow deformation of the Hengelo rock salt, up to strains of about 10 %, with a significant role

played by solution-precipitation processes in controlling rheology.

5.2.5.4 Deformation mechanisms in salt domes Surprisingly few studies are available of the deformation mechanisms and rheology of rock salt under conditions of natural salt tectonic deformation. Microstructures of naturally deformed domal rock salt samples (Speisesalz)

Flow and Transport Properties of Salt Rocks

extremely fine-grained (~0.6 mm). Microstructures such as oriented fibrous overgrowths and growth banding (observed in gamma-irradiated sections, figure 5.2.15) suggest that the principal deformation mechanism was solution-precipitation creep accompanied by grain boundary migration and grain boundary sliding. Crystal fabrics measured by EBSD show only a weak crystallographic preferred orientation consistent with solution-precipitation accommodated grain boundary sliding. Using published flow laws for this mechanism (Spiers et al. 1990), plus slope-based estimates of the gravitational shear stress driving glacier flow, the strain rate in the finegrained mylonites was estimated to be about 10-10 s-1. Figure 5.2.14. Typical micrograph of Hengelo Rock salt, decorated to show the microstructure, with subgrains (white lines), grain boundaries (dark bands), showing clear evidence for “overgrowth” due to solution-precipitation processes such as pressure solution and and grain boundary migration. Mean grainsize in Hengelo samples is between 5 and 25 mm. Width of image is 7 mm (from Schléder and Urai 2005)

from the Asse mine in Germany were described by Urai et al. (1987). All samples consistently showed the operation of dislocation creep processes, accompanied by extensive water-assisted grain boundary migration. Grain boundaries were shown to contain brine films during recrystallisation, and solution-precipitation processes were inferred to have been significant deformation mechanisms. In contrast to the weakly deformed samples from Hengelo, relicts of primary grains in this material were not found, probably because of extensive grain boundary migration.

5.2.5.5 Salt glaciers Microstructural processes in mylonitic shear zones from extrusions of Eocene-Oligocene rocksalts from the Garmsar hills and Eyvanekey plateau (central Iran) were recently were studied by Schléder and Urai (2007), using transmitted light microscopy of gamma-irradiated thin sections, subgrain size palaeopiezometry of polished and chemically etched samples and texture measurements by EBSD. The less deformed “protomylonites” found mostly in the stem region of the emerging diapirs comprise 2-6 mm sized grains, occasionally rich in primary fluid inclusions indicative of their primary non-recrystallised state. Abundant, well-developed subgrains suggest that the protomylonite deformed mainly by dislocation processes. Elongated subgrains at grain edges point to recrystallisation by fluid-assisted grain boundary migration. Recrystallised, strain-free grains are common. The material in the highly deformed mylonitic zones found in the salt glaciers is

5.2.6 Discussion and outlook Through integration of all available data, it is now possible to provide a rather complete model of the deformation mechanisms and microstructural evolution of evaporites under a wide range of conditions, including strain rates well below those reached in laboratory experiments. At these low rates, microstructural studies of natural salts and extrapolation, long term in-situ and subsidence measurements, and extrapolation of experimental data for fine grained samples indicate that solution-precipitation creep and fluid assisted dynamic recrystallisation are important processes, significantly contributing to the total strain rate. It is also clear that the rheology of a salt body is much more heterogeneous than previously thought. This explains the common occurrence of m- to km- scale folding in the interior of salt domes but also in weakly deformed flat-lying salt. The rise of salt to the surface through the cold diapir stem leads to high stresses in relatively strong rock, and this in turn enhances dynamic recrystallisation and grain size reduction when the salt is exposed to rainwater at the surface, so that the strong, cold salt in the diapir stem is dramatically weakened and can flow down the slopes of salt mountains. The available data provide a reasonable basis for modelling the mechanical behaviour of salt under geotechnical and natural conditions. However, in many current studies the effects of water-activated grain boundary processes are often neglected, and this omission must lead to errors in prediction of displacement rates, especially over long periods. Geo-mechanical modelling efforts can thus be significantly improved by making full use of the data available on the effects of water, and some of the discrepancies seen in experimental data on different salts can probably also be explained in terms of these effects.

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J.L. Urai · Z. Schléder · C.J. Spiers · P.A. Kukla Figure 5.2.15. Typical microstructure of solution-precipitation deformation in glacier salt from Iran, as observed in gamma-irradiated sections (after Schléder and Urai 2007). Microstructures such as oriented fibrous overgrowths on both sides of a grain boundary, growth banding and the absence of slip lines or subgrains suggest that the principal deformation mechanism was solution-precipitation creep accompanied by grain boundary migration and grain boundary sliding. Crystal fabrics measured by EBSD in these samples show only a weak crystallographic preferred orientation consistent with solution-precipitation accommodated grain boundary sliding. Image width is 4 mm

Figure 5.2.16. Conceptual model of the rock salt cycle and associated rheological transitions. In undisturbed rock salt (1) primary microstructures reflect the deposition conditions with common microstructure of fluid- inclusion-outlined grains. In slightly deformed rock salt subgrains develop, together with incipient grain boundary migration (2). As the grain boundary sweeps through a fluid-inclusion- rich part of a grain, the primary inclusions are collected at the grain boundary. Alternatively, in fine grained salt solution-precipitation creep is dominant. Large contrasts in rheology are common in layered salt, from the onset of tectonic deformation. As the salt rises, and deforms further (3) steady-state microstructure is produced after complete recrystallisation, although primary fluid-inclusions are still preserved. In the cold diapir stem high stress and small subgrains develop (4). At the surface (5) weak mylonitic shear zones form allowing the salt glacier to flow downhill at unusually high strain rates (6). Finally, dissolution of the salt glaciers in rainwater and redeposition produces secondary evaporites

Additional improvements can be made by obtaining a more detailed understanding of the mechanism of both transient and steady state dislocation creep, and by improving microphysical models for the effects of solutionprecipitation creep, recrystallisation and surface energy driven grain boundary and crack healing on flow and transport properties.

Further work is also needed on deformation mechanisms in naturally deformed rock salt from a wider range of geological settings, and in comparing constitutive equations obtained in the laboratory with those obtained from inverting surface displacement data obtained above salt mining sites and salt extrusions.

5

Chapter 5.3

Dynamics of salt structures P.A. Kukla · J.L. Urai · M. Mohr

5.3.1 Introduction Salt tectonics play a major role in the development of many sedimentary basins. One classic area of salt tectonics is the Central European Basin System (CEBS). Here, the mobile Permian Zechstein salt formed a large number of salt structures such as anticlines, diapirs, pillows, sheets, stocks, and walls (Fig. 5.3.1) during an extended period of salt tectonic activity in Mesozoic and Cenozoic times (Kockel 1998; Jaritz 1973). Major changes in sedi-

mentation patterns and structural regimes are associated and common in this setting (Kockel 2002, 2003). Over the past two decades, however, the focus of salt tectonics studies has been outside this area, being mainly initiated by hydrocarbon exploration in major offshore salt basins worldwide such as the Gulf of Mexico and the West African and South American Atlantic Margins. In Germany, salt dynamics‘ research started to evolve at the beginning of the last century with concepts such as

Figure 5.3.1. Location of the central part of the Southern Permian Basin and the distribution of salt diapirs and pillows (after Lokhorst et al. 1998). The basin margin is marked by the facies changes of the Zechstein 2 carbonates from slope (grey) to basin (light grey). The study area in East Frisia is demarcated by the rectangle

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buoyancy (Arrhenius and Lachmann 1912) or tectonics (Stille 1910, 1925) being proposed as the driving forces for salt movements. The buoyancy-driven halokinetic model of Trusheim (1957, 1960) who postulated an autonomous, isostatic rise of salt and piercement of the overburden due to Rayleigh-Taylor instabilities, was a breakthrough in research and defined the way of thinking regarding salt tectonics in the CEBS for the next forty years. Supplemented by the theory of prograding gravity instabilities forming successive generations of diapirs and salt dome families (Sannemann 1968), this concept of gravity inversion was later modified by assuming a possible tectonic reason for the initiation of salt movement (Meinhold and Reinhardt 1967; Rühberg 1976; Jaritz 1973, 1987, 1992; Zirngast 1996). Brink et al. (1992) proposed a strong connection between diapirism and strike-slip faulting, including rim-syncline rotation. Ge et al. (1997) interpreted diapiric families and peripheral sinks in North Germany as being the result of prograding sedimentary wedges and lateral migration of salt. Recent studies exclude density inversion and suggest decoupling of salt and structurally triggered salt movement as an initiating mechanism in different parts of the CEBS (Mohr et al. 2005; Maystrenko et al. 2005b; Scheck et al. 2003a; Kossow and Krawczyk 2002). Baldschuhn et al. (2001) also argue for an important structural control during several tectonic phases proposed by Kockel (1998). Studies from adjacent areas such as the Southern North Sea explain salt pillows as buckle-folds due to inversion and compression (Coward and Stewart 1995) or by downbuilding (Buchanan et al. 1996). Others propose basement-involved extension and compression as the reason for salt flow (Remmelts 1995) or document raft-tectonics as an extreme mode of extension (Thieme and Rockenbauch 2001) for initiating salt movement. The role of sediments influencing diapir growth by differential loading is an accepted major mechanism for salt movement in other salt provinces (e.g., Jackson and Talbot 1991). Summarising the previous studies, it is now widely recognised that the CEBS was subjected to a multi-phase tectono-sedimentary evolution involving strike-slip, extensional and compressional phases caused by North Atlantic break up and Laramide style inversion in the course of alpine deformation. A discussion of the quantification of salt tectonic processes and their resulting structural geometries and the relationship with the evolution of sedimentary depositional systems during long periods of ongoing basin evolution and changing climates is the purpose of this chapter. It will be demonstrated that an improved understanding of the dynamics and evolution of salt structures within the complexly deformed CEBS can be achieved by an integrated approach to subsurface analysis involving sedi-

mentary and structural analysis based on seismic interpretation, seismic attribute analysis and subsurface data. Initially an overview of the current salt tectonics concepts will be given followed by a detailed case study of an area in NW-Germany in which, based on a 3D seismic dataset, balanced structural retro-deformation and sedimentary analysis were integrated for a better understanding of salt tectonic processes and their corresponding sedimentary response. The chapter will conclude with regional implications for the entire CEBS.

5.3.2 Concepts of salt tectonics The past two decades have seen enormous progress in the understanding of the mechanisms of salt tectonics. This has been summarised in a number of review papers, most recently by Hudec and Jackson (2007) but also Jackson (1995), and Jackson and Talbot (1994). Early salt tectonics workers considered salt movements to be set off by extension or compression (Stille 1925) or passive diapirism (Barton 1933). The prevailing view thereafter was influenced by the work of Trusheim (1957, 1960) in the CEBS who postulated that the mechanics of salt flow were dominated by gravity inversion and halokinesis. In this salt buoyancy model, the yield strengths of sediments had to be disregarded in order to enable a buoyant rise of salt to form a diapir. Subsequent studies in the late 1980s began to recognise the importance of overburden strength as a control on diapir growth. Modern interpretations of salt tectonics emphasise differential loading as the dominant force driving salt flow (Hudec and Jackson 2007). The concept of differential loading and related salt movement explains lateral variation in thickness, density or strength of supra-salt caused by sedimentation or erosion (Jackson 1995; Ge et al. 1997). Sediment wedges prograding and spreading basinward generate lateral pressure gradients in the salt and may induce large scale lateral flow of salt (Humphris 1979). Buoyancy can still be important in a weak overburden setting where the salt is less dense than the surrounding sediments, leading to a continuous rise of salt to the surface. Processes of diapir growth include passive diapirism, differential loading, extension, compression, strike-slip faulting and salt-sheet emplacement. Downbuilding, also called passive diapirism, is a major mechanism for salt movement (Barton 1933; Jackson and Talbot 1991; Vendeville and Jackson 1992a; Buchanan et al. 1996; Rowan et al. 2003). This special type of differential loading involves deposition while salt rises continually with respect to the surrounding strata, while always remaining close to the surface. Downbuilding thus does involve active piercement of the overburden. Sedimen-

Dynamics of salt structures

tary strata often lap onto the flank of the exposed diapir and upturned beds and unconformities along the flank of the diapir constitute halokinetic sequences (Giles and Lawton 2002; Rowan et al. 2003). The geometry of the overall “Christmas-tree” shape of these diapirs largely depends on the rate of salt rise versus sediment deposition rate. Passive diapirism ceases only when salt migration cannot keep up with sedimentation, either because of increasing sedimentation rate or when the evacuation of salt forms a salt weld (Rowan et al. 2003). It is most likely that most of the world‘s tall salt domes and walls spent most of their histories as passive diapirs which may occur in any tectonic setting (Hudec and Jackson 2007). Faulting in the sub-salt sequences usually results in delocalised faulting of the supra-salt, accompanied by lateral flow of salt (Koyi et al. 1993; Nalpas and Brun 1993; Remmelts 1995). Extension is usually a trigger for this type of basement involved reactive diapirism (Jackson and Vendeville 1994). Active piercement is a forceful intrusion of the overburden by a diapir (Schultz-Ela et al. 1993; Vendeville and Jackson 1992a). It only occurs when a critical ratio between diapir height and overburden thickness is exceeded. Another mechanism is extension during thin-skinned gravity spreading and lateral movement of the overburden with no basement involvement. Characteristic features are listric normal growth faults and synkinematic mini-basins (e.g., Gulf of Mexico; Jackson and Talbot 1994). Raft tectonics are an extreme mode of thin-skinned extension, where syndepositional grabens open and the overburden separates into rafts which move downslope (Burollet 1975; Duval et al. 1992; Best 1996; Thieme and Rockenbauch 2001). After the extensional rise of a diapir, trough-like, typically asymmetric, depocentres fill the spaces between the rafts (Vendeville and Jackson 1992b). The unequal extension of basement and overburden is in agreement with the drifting stage of a passive margin (e.g., Kwanza Basin of Angola), but in encratonic basins two combined stress regimes (extensional up-slope and a compressional down-slope) are often developed (Letouzey et al. 1995; Penge et al. 1999). Compression also plays an important but still under-recognised role in salt tectonics (Stille 1925; Koyi 1988; Coward and Stewart 1995; Letouzey et al. 1995; Waltham 1997). The typical structural environments are thin-skinned fold and thrust belts, thick-skinned intracratonic inverted-basins, and shortening at the toe of gravity-gliding systems (Letouzey et al. 1995). Salt-cored folds are often the locus for diapir formation in these settings. Buckle folding in a compressional regime is seen as one mechanism for salt pillow evolution in the North Sea (Coward and Stewart 1995) and the Zagros Mountain Belt (Koyi 1988).

Salt may also be emplaced into its overburden in the hanging wall of a thrust fault or through released bends created by strike-slip movement which causes drastic thinning of the overburden. Raft tectonics occur during extreme thin-skinned extension over a decollement layer (Burollet 1975; Duval et al. 1992; Best 1996; Thieme and Rockenbauch 2001). Rafts are typically separated by trough-like, typically asymmetric, depocenters of younger, syn-kinematic strata (Vendeville and Jackson 1992b). Further geometric features in salt terrains are related to allochthonous salt. This comprises mostly sheet-like salt bodies such as salt tongues, salt stocks (which can be amalgamated as salt-stock canopies) and salt-nappe systems (Warren 2006). Common to all these concepts is the emplacement above the main evaporite salt layer.

5.3.3 Salt geometries and kinematics – a case study 5.3.3.1 Subsurface geometries from seismic interpretation The investigated example area is located at the SW margin of the CEBS and the western flank of the Triassic Ems Graben (Fig. 5.3.1) in the proximity to major basin faults. The data used for this study consists of pre-stack depth migrated 2D and 3D seismic data from a 30 x 30 km area, including deep boreholes and a large number of non-depth migrated 2D seismic sections with a total length of approximately 430 km. In general the seismic data images the sub-salt basement very well. Nevertheless, typical difficulties of seismic interpretation such as uncertainties in sub-salt imaging at the flanks of the diapir are also apparent in this area. In this study 16 seismic reflectors and a consistent fault network were interpreted in the depth migrated 2D sections and the 3D seismic cube. The sub-salt Permian Rotliegend Formation and its structural configuration is, as an exploration target, the best analysed geological horizon in the area. The pattern of sub-salt Rotliegend faults was compared with the structures in the overburden and the regional geological framework. A typical regional W-E trending seismic transect (Fig. 5.3.2) shows Top Rotliegend rising 780 m from east to west towards the “Groningen High” (Fig. 5.3.1). Faulting of upper Rotliegend age is shown in antithetic and conjugate normal fault sets throughout the area. The Lower and Middle Buntsandstein (SU, SM) sequences in this section show a sedimentary thickness of ~ 650 m. Normal faults in these sequences have a mod-

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Figure 5.3.2. Regional west-east seismic section (depth-migrated) and geological interpretation. The top of the section is the present-day surface. Black arrows in the geological section depict the unconformities mentioned in the text for each section respectively. The section shows the westward rise of basement, the disconnected Lower / Middle Buntsandstein (SU/SM) sequences, three salt structure geometries of the Zechstein (Z) and a large depocentre of the Lower Keuper (KU) in the western central part. For further details see text

erate easterly dip in the west and a westerly dip in the central part. In the central and western part of the profile there are two areas close to larger salt structures where Lower and Middle Buntsandstein sequences are either absent or developed as “rafts” within salt lithologies. The Upper Buntsandstein (SO) thickness varies between 500 m in the west and 100 m in the centre of the area. Three disconnected blocks are interpreted, located directly above Zechstein salt. A former connection of these blocks is indicated by good correlation of the seismic reflectors across their broken edges. The disconnected Upper Buntsandstein blocks are covered and surrounded by displaced Zechstein salt.

rests unconformably on Upper Muschelkalk. In the central part of the section, the interpreted Middle Keuper to Early Jurassic rim-syncline belongs to the salt diapir A. The base Cretaceous unconformity truncates the Triassic sequences down to the Muschelkalk in the west and down to the Jurassic in the east (Fig. 5.3.2). Cretaceous to Quaternary sequences dip westward and increase their thickness along equally westward dipping normal faults. The salt structure in the western part of the profile shows large variations in thickness of Zechstein salt below Upper and Lower/Middle Muschelkalk sequences and its top truncating sequences up to at least Middle Keuper level.

The Lower and Middle Muschelkalk (MU, MM) show a large thickness increase from 250 m to 650 m (east to west) and are missing in the western part of the section. The lower boundary of the Upper Muschelkalk Formation is formed by an unconformity in the central part of the profile.

In comparison, a N-S trending longitudinal section in figure 5.3.3 is positioned at low angle to the regional structural grain. Southward dipping normal faults in the sub-salt basement show a maximum heave of 250 m. The Lower/Middle Buntsandstein is faulted, truncated and partially eroded in the central part of the profile. In the northern part these sequences are absent and Upper Buntsandstein (SO) lies directly above Zechstein salt. In the south, blocks of uppermost Buntsandstein and Lower/Middle Muschelkalk are domino faulted, tilted and displaced, most probably on Upper Buntsandstein Röt salt. These faulted blocks in the south and the flatlying sequences of Lower/Middle Muschelkalk in the north are truncated by an unconformity at the base Up-

Three normal faults cutting this sequence are spatially associated with sedimentation of a local depocentre in the Lower Keuper (MU) and lowermost Middle Keuper (Grabfeld-Formation, KM1) with a maximum thickness of 1,000 m. Their seismic reflectors converge eastward to diapir A. The sequences are absent adjacent to the diapir, where the Stuttgart-Formation (Middle Keuper 2, KM2)

Dynamics of salt structures Figure 5.3.3. Regional north-south seismic section (depth-migrated) and geological interpretation. The section shows absence of the Lower/Middle Buntsandstein sequences in the north. Faulted and tilted blocks of Upper Buntsandstein and Lower/Middle Muschelkalk are truncated by an unconformity at the base of Upper Muschelkalk. For further details see text

per Muschelkalk (MO). Upper Muschelkalk sequences are thickened in the central part, where they rest directly upon Zechstein salt. The overlying Lower Keuper Formation has its maximum depocentre in the northern part of the profile, where Lower/Middle Buntsandstein sequences are absent or incomplete. In addition, the lower Keuper shows syn-sedimentary normal faulting and

southward converging seismic reflectors. The base Cretaceous unconformity truncates the Triassic sequences down to Middle Keuper (KM4) level in this section but down to the Muschelkalk in others. Cretaceous to Quaternary sequences dip slightly to the south-west with an increasing thickness and are intersected by westward dipping normal faults.

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Figure 5.3.4. Sequential retro-deformation (structural balancing) of a 30 km long west-east seismic section in the study area. Three phases of salt movement can be identified. During the first phase (Zechstein to Middle Triassic) thin-skinned extension and rafting is accompanied by lateral salt movement and the growth of a salt pillow in the central part. The second phase (Late Triassic to Jurassic) is defined by extension which triggered diapirism followed by sedimentary downbuilding and periods of diapir emergence. The last and third phase of salt movement was triggered by Late Cretaceous to Tertiary compression that caused the growth of buried salt structures. For further details see text

Dynamics of salt structures

5.3.3.2 Salt tectonic evolution based on retro-deformation Several authors (e.g., Rowan 1993; Hossack 1995; Buchanan et al. 1996; Schäfer et al. 1998) have demonstrated that palinspastic restoration in salt basins is a useful tool to reconstruct the structural and sedimentary evolution over time. Strata and structures are removed sequentially to unravel the kinematic evolution and the corresponding salt geometries. In addition, different subsurface geometries discerned during seismic interpretation can be tested and uncertainties constrained. Since seismic interpretation has often non-unique results, an iterative process of structural balancing followed by an update of the interpretation should be used. The final model is consistent with all known constraints and is therefore more likely than others, but not necessarily unique (Mohr et al. 2005).

below as a consequence of salt movement and sedimentary response (Hossack 1995). The usual restoration algorithms in salt regimes are oblique shear and flexural slip. Oblique shear preserves area, but line-length is not constant, which causes a significant length loss when restoring steeply-dipping layers. Most workers prefer oblique shear for downbuilding and extension while flexural slip is preferred for structural shortening and active salt diapirism. Fault displacement may be removed by fault-parallel flow, inclined shear or using move and rotate tools. Sequential retro-deformation modelling was used to unravel the tectonic history of this part of the CEBS. A series of balancing iterations were adopted in cases where conflicts with geological data or model requirements were found. Figure 5.3.4 is based on the regional seismic section of figure 5.3.2, restored to 12 important time slices. The dynamics of this part of the CEBS can be broken down into three major phases which will be described sequentially in the following section.

In salt structure restoration the overburden evolution is seen as a direct result of salt flow. Salt itself can only be passively restored. In addition, salt area may vary due to dissolution or out-of-section migration and the algorithms developed for retro-deformation of brittle rocks cannot reproduce the kinematics of the salt (Rowan 1993). In a full restoration, the effects of sedimentation, compaction, isostasy, thermal subsidence, and the dynamics of faulting and salt movement should be included. Structural restoration models shown in the following contain sedimentation, fault kinematics and salt movement as the main factors. Variations in isostasy and thermal subsidence are negligible for typical regional balanced sections with a length of tens of kilometres. Compaction influences the vertical thickness of sedimentary sequences over time, but is only relevant when lateral variations in thickness or facies are present, then affecting mainly the uppermost sequences. Sensitivity analyses have shown that disregarding decompaction does not influence the results for the mechanisms and processes of structural and salt tectonic evolution interpreted from section restoration. In cases like the following, where the primary focus lies on the mechanisms and processes rather than the rates and masses, decompaction of sections is not a requirement.

In this model, decoupled basement extension at the basin margin is regarded as the trigger for initial salt movement as is seen in other such basins (Jackson and Vendeville 1994; Vendeville and Jackson 1992a). This scenario includes detached faulting of the cover sequences and continued extension which initiated salt diapirism and the development of sedimentary “rafts”. Salt flows towards the extensional structures in the supra salt, in particular above the basement fault and towards the distal zone of detached extension (Koyi et al. 1993). Between these structures differential sedimentary loading creates an increased sedimentary thickness until a salt weld is formed and lateral salt flow stops. As a consequence, the diapir falls and a mini-basin grows in the intra-raft area (Vendeville and Jackson 1992a,b). Salt migrates laterally along dip and forms an adjacent pillow-like salt structure basinward.

An essential assumption in 2D structural balancing is plane strain, and that the section should be orientated parallel to the tectonic transport direction. The decoupling by salt necessitates the treatment of the salt layer, the sub-salt and supra-salt sequence as three autonomous tectonic systems during the restoration process (Schäfer et al. 1998). In addition, the positioning of the regional elevation or target horizon to which a template line is restored is of importance. Regional elevation represents the pre-deformational relief, which is deduced from a point without deformation or a line that defines area balance above and

In the investigated area, we observed a major spatial gap between Lower and Middle Buntsandstein rafted blocks in the vicinity of salt bodies and with a significantly thinned section compared to the coherent sections in the west and east (Fig. 5.3.2). Seismic reflectors at the edges of the Lower/Middle Buntsandstein blocks are truncated by faulting and several faulted blocks lie isolated in the centre of the gap. In the sub-salt basement of the working area in general, no corresponding normal faults intersecting rafts have been detected. These features of thin-skinned extension are consistent with the regional

5.3.3.3 First phase of salt movement: Zechstein to Middle Keuper

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Figure 5.3.5. Vertical thickness analysis from a seismic 3D dataset allows evaluation of structural trends, regional stress and depositional patterns through time. Thickness maxima are shown in red color, thickness minima in light brown color. Maps reveal changes of structural patterns and depocentre migration between geodynamic phases (compare with figure 5.3.4) in the CEBS. For further details see text

Dynamics of salt structures

subsidence trend to the east (Röhling 1991). In addition to rafting at the northwestern margin of the Ems Trough described here, coeval rafting events are known from other marginal positions along the basin (Kockel 2002). Outside the working area an Early Triassic NE-SW directed basement step at the western basin margin (Baldschuhn et al. 1991) is documented. Basement faulting caused salt diapirism directly above (Baldschuhn et al. 1991) as well as detached supra-salt faulting and increased Middle Buntsandstein to Middle Muschelkalk layer thickness in the hangingwall - both visible in our data (Fig. 5.3.2). Two extensional phases are documented by the rafted blocks of Lower/Middle Buntsandstein and Upper Buntsandstein. The sedimentary thickness distribution shows a southeastward migration of the Upper Muschelkalk to lowermost Middle Keuper (KM1) depocentres (Fig 5.3.5g-i) onto Zechstein salt and normal regional thickness above the former depocentres in the west (Fig. 5.3.2). The primary salt-source layer in the northwest was probably exhausted, resulting in the fall of the diapir and the development of a mini-basin with increased sedimentary thickness in the intra-raft area (Fig. 5.3.5i). The Westdorf Graben in the northern continuation of the studied zone, contains an extreme thickness of Upper Buntsandstein to Lower Keuper sediments (Gaertner and Röhling 1993), which corresponds to a zone of absent Lower/Middle Buntsandstein sequences (Röhling 1991). During detailed retro-deformation Top Zechstein level (Fig. 5.3.4l) is modelled with ~ 0.3° eastward dip and Top Rotliegend starts without faults and with minor relief dipping ~ 0.6° to the east. We assume an original thickness of the Zechstein salt of between 820 m and 940 m increasing to the east based on our modelling results. The Lower and Middle Buntsandstein sequences (Fig. 5.3.4k) are interpreted to initially completely cover the salt with an eastward thickness increase from 540 m to 760 m corresponding to the regional dip. The light pink area demarcating missing Lower/Middle Buntsandstein at the eastern end demonstrates the difference between section length and sequence length, and thus a regional extension between Top Middle Buntsandstein times and present day (Fig. 5.3.4k-i). Before Upper Buntsandstein times (Fig. 5.3.4j) the first extensional event caused ~ 950 m (~ 3%) extension in the supra-salt by faulting and rafting, whereas the sub-salt is unfaulted. Differential extension at section scale caused by structural decoupling must be balanced in the complete intra-continental basin (Letouzey et al. 1995). Basement faulting elsewhere at the basin margin and shortening in the basin centre can compensate for this thin-skinned extension. The early stage of extension and

rafting initiated reactive salt diapirism, followed by active piercement of thinned overburden by salt and possible extrusion at the surface. We propose that salt may have reached the surface in both events, but was covered soon thereafter by younger sediments (Fig. 5.3.4k-j). Sediments of the Upper Buntsandstein covered this early salt diapir (Fig. 5.3.4j) and show a laterally variable thickness distribution with increased sedimentary thickness to the west and a decreased thickness adjacent to the intra-raft due to differential loading and lateral basinward salt flow (Fig. 5.3.5h). A slight thickness increase to the eastern part of the section is interpreted as the result of regional dip to the east. The Lower / Middle Muschelkalk shows a slight westward migration of its depocentre in comparison to the underlying sequence. The primary salt source at this stage is nearly exhausted in the western part of the section. Models and data suggest another extensional event that has also triggered rafting and diapirism before the beginning of Upper Muschelkalk times. Extension calculated from the balancing for this phase was ~ 1650 m (~ 5.4%). Salt rise is accompanied by drag folding and erosion at the flanks of the Middle Muschelkalk diapir. An unconformity truncated this sequence before the beginning of Upper Muschelkalk sedimentation that covered the former diapir. Reconstruction to the Top Grabfeld-Formation (Fig. 5.3.4i) shows major mini-basin growth of ~ 1100 m above the intra-raft area and eastward adjacent ~ 300 m erosion down to the Top Middle Buntsandstein sequences. Salt flow from the west then stopped, initiating the fall of the diapir due to the eastward regional dip. Differential loading, sedimentary prograding and lateral salt flow resulted in the formation of the Lower Keuper depocentre, the pillow-like salt structure and the primary rim-syncline in the eastern end of the section. Erosion and faulting at the crestal zone of the salt anticline weakened its roof and facilitated the collapse of the structure.

5.3.3.4 Second phase of salt movement: Middle Keuper to Top Jurassic The second phase of salt movement is interpreted as having started as the consequence of sub-salt extension that triggered normal faulting in the overburden, as described in the concept of reactive diapirism (Vendeville and Jackson 1992a; Jackson and Vendeville 1994). Reactive diapirism followed by a short phase of active diapirism passes into passive diapirism (Vendeville and Jackson 1992a), also known as downbuilding (Barton 1933; Jackson and Talbot 1991; Vendeville and Jackson 1991, 1992a; Buchanan

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et al. 1996). Typical features are near-diapir onlapping, salt re-sedimentation, increased thickness, unconformities and salt jags at the edge of the down-built sedimentary wedges (Giles and Lawton 2002; Rowan et al. 2003). This “Early Cimmerian” extensional phase was caused by intensified North Sea rifting that brought about major extension and faulting in the onshore Ems region or in the Glückstadt Graben (Ziegler 1990; Frisch and Kockel 1999). Strong dilation was associated with the beginning of major salt diapirism, subsequent rim-syncline development and extrusion of Zechstein salt in this and many other parts of the North German Basin (Frisch and Kockel 1999; Baldschuhn et al. 2001; Jaritz 1973). As a result of this tectonic reorganisation, it can be observed in the study area that during this phase depocentres above the major unconformity at the base of the StuttgartFormation (KM2) migrated to the flanks of the present diapirs (Fig. 5.3.5d-f). Since then, the sedimentary thickness distribution is directed along a NNW-SSE axes parallel to diapir A and B. The underlying older suprasalt sequences in contrast show erosional surfaces at the flanks of these diapirs and a NNE-SSW to NE-SW directed thickness distribution and faulting oblique to the latter structures (Fig. 5.3.5d-f). A striking change in structural patterns is therefore inferred for the end of Grabfeld Formation (KM1) deposition. Major faulting in the sub-salt basement, seen as the reason for this evolution, formed a graben structure that corresponds to this NNW-SSE structural direction of the supra-salt. Seismic indications for the downbuilding phase are Middle to Upper Keuper sedimentary wedges with basal unconformities and onlap structures as well as salt jags of the laterally extruded diapirs (see below). Jurassic sequences onlapping the diapirs indicate slower salt rise probably caused by exhaustion of the salt source layer. The base Cretaceous unconformity truncated the Jurassic at the flanks of the diapirs and the Muschelkalk sequences in the west. This is in good agreement with the results for uplift, tilting and erosion of this area in the context of the Late Jurassic doming event in the Southern North Sea (Ziegler 1990). At the end of Grabfeld Formation deposition (Fig. 5.3.4h-i) the basement underwent extension by normal faulting and formation of an asymmetric graben structure with major bounding faults in the west. The amount of basement extension inferred from balancing was ~260 m (0.9%). Basement extension triggered normal faulting and extension in the overburden with an elongation of ~1120 m (3.7%). Supra-salt extension is focussed in two zones. One site is in the west in the vicinity of a major basement fault. It shows minor salt thickness and is located above

the previously rafted area. The other zone of focussed extension had already been weakened by previous faulting and erosion above the central salt structure. Salt pierced the overlying sequences as a reaction to extension and extruded to the surface, starting a downbuilding process with peripheral sedimentary sinks as secondary rim synclines developing. The reconstructions to the Top Weser Formation and to Top Keuper times (Fig. 5.3.4g, f) show the central diapir with continued downbuilding and salt at the surface. The migration of the local Middle Keuper subsidencemaxima towards the diapir (Fig. 5.3.5f) indicates the successive collapse of the former salt pillow. The almost constant sedimentary thicknesses of the sequences west of the diapir suggests less salt movement on this side (Fig. 5.3.5d-f). We suggest that outside the graben structure only a thin section of Middle and Upper Keuper (~120 m) is preserved in the west and none in the east, in contrast to the area in between (up to 1,600 m). The western salt structure could not evolve as a diapir because of the exhausted salt source and it is therefore modelled to be already covered by the Arnstadt-Formation. At the eastern end of the profile a pillow-like salt structure evolved above a westward dipping basement fault after the deposition of the Weser-Formation (Fig. 5.3.4h). Normal faulting after sedimentation of the Weser-Formation resulted in reactive diapirism and subsequent downbuilding at the easternmost salt structure (Fig. 5.3.4f). Ongoing basement extension through the Middle Keuper ceased before the Upper Keuper. For restoration to the Top Jurassic time slice (Fig. 5.3.4e), eroded Muschelkalk to Keuper sequences and a thickness of 550–700 m of Jurassic deposits have been assumed. The sub-salt basement is modelled with 1.1° regional dip. The termination of downbuilding is proposed for the Middle to Late Jurassic, when salt was completely removed from beneath the sedimentary pile, salt rise ceased and the diapir was covered by sediment (Fig. 5.3.4e). Before the beginning of sedimentation in the Early Cretaceous, more precisely the base late Hauterivian (~128 Ma), the last tectonic pulse of the “late Cimmerian” phase took place. We modelled this by regional uplift and tilting. Tectonic uplift of 1,170 m in the west and 650 m in the east produces a ~2.1° eastward-dipping sub-salt basement. This enabled erosion of most Jurassic sequences, parts of the Keuper and Muschelkalk and the uppermost ~160 m of the salt diapirs roof. Using seismics, we interpreted two disconnected Buntsandstein blocks in the Central salt structure, generated by supra-salt extension and embedded in Zechstein

Dynamics of salt structures

salt (Fig. 5.3.4e-h). Kinematic and geometrical modelling provide no further constraints on whether these blocks sink into the salt or even ascend with the rising salt. Using simple mechanical considerations based on Stoke’s law, one can estimate the rate of sinking of a Buntsandstein block in salt. For example, we calculate the velocity for the westerly block using 250 m for the radius and 2,500 kg/m3 for the density of the Buntsandstein. We used 2,160 kg/m3 as an average value for the density of salt. The critical parameter in the calculation is the salt viscocity which may vary between 1017 Pa s for fine grained salt at relatively high temperatures and 1020 Pa s for coarser grained salt at relatively low temperatures (Van Keken et al. 1993). This results in a steady-state velocity of 15 m/Ma to 15*103 m/Ma. The salt’s velocity when rising is estimated at between 1,2*103 m/Ma and 1,0*102 m/Ma for the different time steps. These simple calculations indicate that the Buntsandstein blocks may or may not sink. For a more accurate constraint, more precise data on halite rheology are required.

5.3.3.5 Third phase of salt movement: Early Cretaceous to recent The process dominating this latest phase of salt movement is an episodic horizontal shortening of the salt structures (cf. Vendeville 2002). Regional compression rejuvenates diapiric growth by squeezing the salt structure which is weaker than the higher-strength sediments laterally adjacent. Deformation is concentrated in the salt structure and in the sedimentary sequences above. The latter should be affected by regional shortening and additionally by diapiric rise. In this model, structural decoupling is required between sub- and supra-salt, but no additional salt source is needed to rejuvenate diapiric growth. Support from seismic data for this kind of model in the investigated CEBS comes from the positive relief of the salt structures in relation to the regional datum indicating Cretaceous to Cenozoic uplift. Peripheral sinks are weakly developed, so that the volume of rising salt can not be balanced by the volume of sediment in the peripheral sinks. Sedimentary sequences above the salt structures show erosional unconformity surfaces, toplaps and reduced thickness in distinctive layers. Pairs of conjugate faults observed on top of the diapirs can be interpreted as the result of oblique compression (Fig. 5.3.5b,c). Reverse faults, minor folding and arching of layers as well as the formation of crestal grabens are interpreted. Support from the regional evolution comes from the different phases of Cretaceous to Cenozoic compressional tectonics under a N-S to NESW stress field that are well known from inversion of German and Dutch basins (Kockel 2003; De Jager 2003). Using an iterative process of re-deformation to Base Tertiary and Base Campanian, we first used vertical shear as

the restoration algorithm for a static and solely vertical movement of salt and sediment. In this case, vertical shear affects the layers producing an unsatisfactory geometry of the diapir and the adjacent sediments. The use of flexural slip to model a small amount of shortening on the other hand reasonably reproduces the previous configuration. The Early Cretaceous sedimentary sequence displays a slight thickness increase to the west and a reduced dip of the sub-salt basement is visible in the restoration to the base Campanian (Fig. 5.3.4c). This trend is observable throughout the entire Cretaceous (Fig. 5.3.4c, d). Subsurface salt dissolution above areas of subcropping salt of the Middle Muschelkalk or Middle Keuper caused westward dipping normal faults and a local thickness increase of the Cretaceous units. Differentiated sedimentary thicknesses at the top of the diapirs (Fig. 5.3.4c) are also interpreted as a subrosion effect. Layer thicknesses of the uppermost Cretaceous and Palaeogene (Fig. 5.3.4a, b) are clearly thinned at the roof of the central diapir, accompanied by only a minor thickness increase at the flanks (Fig. 5.3.5c). This unbalanced diapiric rise of ~400 m above the regional datum is modelled as a thin skinned shortening event (Fig. 5.3.4b, c) caused by Late Cretaceous and Tertiary compression. Strain is concentrated in the weak and incompressible salt of the squeezed and uplifted diapir. Calculated shortening from modelling is around 0.3% (~100 m). We identified two phases of shortening in the uppermost Cretaceous and in the Palaeogene that ended before the Neogene (Fig. 5.3.5a,b). The present-day structure (Fig. 5.3.5a) as the source for restoration includes a ~1.5° eastward dipping sub-salt basement and a uniform Neogene layer thickness. There are no indications for recent salt movements in the area.

5.3.4 Salt sediment interaction In salt basins, sedimentation processes vary regionally in response to changes in sea level and climate whilst salt tectonism plays an increasing role during burial of such basins (Vendeville and Jackson 1991; Talbot 1995; Giles and Lawton 2002; Rowan et al. 2003; Schultz-Ela 2003). In the CEBS, salt-influenced sedimentary responses to renewed phases of tectonism can be clearly discerned from detailed sequence analysis based on seismic and log data combined with retrodeformation modelling studies. Late Palaeozoic sedimentation in the CEBS deposited Upper Rotliegend sediments in a series of playa mudflat – sandflat settings in an extensional regime. About 800 m of bedded sulfate and halite were deposited in the study

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P.A. Kukla · J.L. Urai · M. Mohr Figure 5.3.6. Detailed tectono-sedimentary analysis at the SW flank of a salt diapir based on depth-migrated 2D seismic section (top). Geological interpretation (base) shows four sedimentary wedges of Keuper age and the salt flanges of diapir “C”. The wedges of Middle Keuper 1 and 2/3 show lowangle unconformities and onlaps on the narrowing diapir. The salt flanges were formed during constant salt rise and extrusion on the land surface with no sediment accumulation. In a sequence stratigraphic context, the unconformities at the base of the sedimentary wedges represent erosional event sequence boundaries. See text for further details

area during the hydrographic isolation and drawdown of the Late Permian Zechstein evaporite basin (Jaritz 1973; Warren 2006). Triassic sedimentation changed from continental clastics during the Early Triassic Buntsandstein to shallow marine sequences of carbonates and evaporites in the Middle Triassic Muschelkalk and back into continental clastics with intercalated saline-pan evaporites during Late Triassic Keuper time (Fig. 5.3.6). Several phases of Triassic rifting influenced the area and triggered its multiphase salt tectonics. Kinematic restoration modelling above showed that decoupled extension in the late Early (Middle Buntsandstein) and Middle Triassic (Middle Muschelkalk) initiated salt diapirism and lateral salt flow with growth of a pillow-like salt structure. Subsequent formation of a basement graben in the early Late Triassic triggered reactive diapirism and the collapse of the salt pillow followed by passive diapirism during which salt remained at or near the surface with the potential of feeding peripheral salt glaciers. The extrusion process stopped in the latest Late Triassic and Jurassic as the falling diapir crest was covered by marine sediments (Mohr et al. 2005). A major unconformity at the Base Cretaceous marks a subsequent episode of regional uplift and tilting in the Middle Jurassic to Early Cretaceous, accompanied by erosion and salt dissolution (Fig. 5.3.6). Late Cretaceous

to Early Tertiary compressional tectonics renewed vertical salt movement under substantial sedimentary cover (Mohr et al. 2005). Triassic Keuper salt glaciers (namakiers) Seismic variance analysis of the high resolution 3D seismic data set clearly images (buried) salt extrusions ~50 m below the Base Cretaceous unconformity, presently at ~2 km depth (Mohr et al. 2007). The variance map shows the morphology of the buried salt glacier to be a fan-like salt apron that covers older sedimentary layers (Fig. 5.3.7). The tongue of the buried extrusion has a jagged outline and an irregular rugged surface with a divergent lineation from east to west (Fig. 5.3.7A, B). This contrasts with the smooth surfaces of the adjacent clastic sedimentary strata. The extrusion is on the western flank of a NNE-trending salt diapir, which has a high variance contrast. For detailed interpretation of the variance map and for stratigraphic calibration we used additional information from the 3D seismic data volume and wireline data from four nearby exploration wells (Fig. 5.3.7B). Borehole A is located 500 m to the southwest of the diapir which has a notable bulge in its westernmost part. West of the bulge a surrounding rim syncline is not developed, whereas elsewhere it is typically filled with Upper Keuper to Jurassic sediments. The salt glacier, probably the residue of a former

Dynamics of salt structures

Figure 5.3.7. Subsurface salt glacier (namakier) of Middle Keuper age. A) Variance map 50 m below the unconformity at the Base Cretaceous displays the uppermost salt extrusion as a fan like structure inside the circle. B) Geological interpretation of A showing the geometry of a Late Triassic salt glacier, the position of the two cross sections and the location of the wells. C and D) Enlarged seismic sections (2.5 x exaggerated) showing four salt glacier generations and the salt diapir, and sedimentary sequences. The vertical dashed lines mark the intersection of the two seismic profiles. The trace of well A is shown by the doted line in the southern part of section D

larger salt extrusion, widens into the direction of flow to the SW. It covers an area of 2.4 km2, is 2.7 km in length and ranges between 0.5 km and 2.2 km in width. The southern end of the structure trends to the SE and is enclosed by a small graben structure. The topographic low of the graben most probably formed the preferred flow path for the namakier. There is also a NE-SW trending low north of the salt extrusion. This depression, controlled by a normal fault (Fig. 5.3.7B), is mostly covered by the namakier. A seismic cross section of the mapped namakier sequence (labeled as salt glaciers I-IV in figure 5.3.7C and D) displays a 50 m to 100 m thick unit of lenticular shape defined by strong seismic reflectors. The strong seismic reflectivity of the glaciers is either caused by a thin salt layer or represents a residual clay and anhydrite layer related to a dissolution carapace. The lenticular cross section of the salt glacier is clearly traceable through the entire 3D seismic volume and can be distinguished from the flanking sedimentary layers. Three previous salt extrusions of Middle Keuper age are identified in seismic sections within the Weser and Arnstadt Formations and indicate classic “Christmas-tree” morphologies (Figs. 5.3.6 and 5.3.7). The seismic profiles show salt wings connected to the westerly diapir which are characterised by irregular low-contrast reflectors that are similar in character to those of the salt mass of the diapir and thus clearly distinguishable from the adjacent sedimentary layering. For example, the lowermost salt glacier (IV) (Fig. 5.3.7 D) forms a wedge of ~500 m thick allochthonous salt between sedimentary strata. This oldest extrusion is only

visible in the N-S section (as seen in figure 5.3.7D) because it spreads to the north, in contrast to salt glaciers II and III, which extend to the northwest and the youngest salt glacier (I) that shows a southwesterly flow direction. A major normal fault cuts the sedimentary layers in the eastern part of the seismic sections (Figs. 5.3.7C, D). On the variance map (Fig. 5.3.7A) the fault is denoted by a horizontal offset of the sedimentary strata on both sides of the salt glacier. This listric extensional fault with a dip to the NW was active throughout the entire Mesozoic. The fault only cuts the salt diapir during upper Middle Keuper to Upper Keuper (kmW-ko; Fig. 5.3.7B). During a near-surface diapiric growth, the variation of the rate of vertical salt rise to the local sediment accumulation rate controlled near-salt sedimentation, salt body geometry, surficial outcropping and extrusion. Typical features are onlapping, drape folding and basal unconformities (Vendeville and Jackson 1991; Talbot 1995; Giles and Lawton 2002; Rowan et al. 2003). Up to the Jurassic, the diapirs in this part of the CEBS (as exemplified by diapir C in Fig. 5.3.6) rose syn-depositionally to a position near the surface, whereby sediment sank into the depleting salt layer in a downbuilding system. The two wedges of Middle Keuper 1 and 2/3 (Fig. 5.3.6) show low-angle unconformities and onlaps on the narrowing diapir. This can be interpreted to be caused by ongoing extension that slows salt flow, narrows the diapir and produces low-angle sedimentary onlaps (Vendeville and Jackson 1992b). In contrast, high-angle onlaps, drape folding and a more vertical diapiric rise are observed in the Middle Keuper 4, which imply the end of Keuper

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extrusion on the land surface with no sediment accumulation. In a sequence stratigraphic context, the unconformities at the base of the sedimentary wedges represent erosional event sequence boundaries generated by sea level low stands (Aigner and Bachmann 1992) or uplift caused by tectonic pulses. The Lowstand Systems Tracts can thus be represented by erosional unconformities (Fig. 5.3.8), salt dissolution or minor sedimentation and salt extrusion of an expanding diapir. Sinking diapir flanks due to salt rise prevent the salt extrusions from eroding. Deposition during the Transgressive Systems Tract causes sedimentary onlapping on a retreating diapir (Fig. 5.3.8). During the Highstand System Tract the diapir can be overstepped by sediment (Fig. 5.3.8). Accommodation space for sediment fill was maintained by the ongoing interplay between regional extension, salt flow and depositional loading in depressions adjacent to the growing diapir.

5.3.5 Multiphase salt dynamics in the CEBS

Figure 5.3.8. Schematic halokinetic sequence stratigraphy of an extrusive diapir. Compare with figure 5.3.6 for sedimentary geometries and positions of sequence boundaries. Unconformities at the base of the sedimentary wedges represent erosional event sequence boundaries generated by sea level low stands or uplift caused by tectonic pulses. The Lowstand Systems Tracts (LST) can be represented by erosional unconformities, salt dissolution or minor sedimentation and salt extrusion of an expanding diapir. Sinking diapir flanks due to salt rise prevent the salt extrusions from eroding. Deposition during the Transgressive Systems Tract (TST) causes sedimentary onlapping on a retreating diapir. During the Highstand System Tract (HST) the diapir can be overstepped by sediment

extension (cf. Frisch and Kockel 1999). The boundaries of the sedimentary wedges widely correspond to the cycles and sequences proposed for the German Keuper basin (Wolburg 1969; Aigner and Bachmann 1992; Frisch and Kockel 1999). Near-diapir sedimentation in the Middle Keuper sequences is dominated by inter-layered salt (Fig. 5.3.6), which can be interpreted as primary evaporites (Frisch and Kockel 1999) preserved due to subsidence or as re-deposited Zechstein salt dissolved and deposited as secondary evaporation products (Trusheim 1971). The salt flanges were formed during constant salt rise and

Results from this case study bear implications for the entire Central European Basin System. The investigated area is situated in the western part of the basin close to major fault systems (Ziegler 1990). Supra-salt extensional faulting which is prominent in the Late Triassic overburden fits well into the regional tectonic setting of the Triassic in the entire basin (Frisch and Kockel 1999; Kockel 2002). Several pulses of basin extension triggered and controlled the main phase of salt diapirism and allochthon growth (Mohr et al. 2005). Regional structure in the study area is characterised by normal faults with adjacent syndepositional depressions that were variably active during the entire Triassic indicating a loading response caused by ongoing detached supra-salt faulting. The initial salt movement in the study area was controlled by extension of the rifting Ems sub-basin at the end of the Triassic Middle Buntsandstein sequence. Basement faulting at the north-western margin of the Ems Trough caused decoupled supra-salt faulting and gravity-gliding along the basin slope. Rafting and initial diapirism was accompanied by differentiated sedimentation. We suggest a broader distribution of these mechanisms in other marginal positions, for example at the edge of the Eichsfeld-Altmark Swell, where the Lower and Middle Buntsandstein are also locally absent. Because rifting in the CEBS started already in lower Middle Buntsandstein (Kockel 2002) this could be the time of basinwide initial salt diapirism and rafting. Locally increased sedimentary thickness since

Dynamics of salt structures

the Upper Buntsandstein points to differential loading and lateral salt flow. After the salt source was exhausted updip of the diapir, salt supply ceased and the fall of the diapir began, accompanied by Lower Keuper mini-basin evolution. Lateral salt-flow produced a pillow-like salt structure at the basinward shoulder of the intra-raft area. We also propose this evolution for the northern continuation of the study area, the Westdorf Graben (Gaertner and Röhling 1993). Assumed Keuper age rafting events (Best 1996; Thieme and Rockenbauch 2001) may also be attributed to this mechanism. Retro-deformation results imply that strong supra-salt extension during thin-skinned tectonism exceeds the amount of extension due to basement faulting, and therefore must be balanced by shortening in the basin centre – associated with reverse faulting and buckle folding. We suggest that this buckle folding is one mechanism for basinwide salt pillow evolution in Buntsandstein and Muschelkalk sequences, especially in cases where no basement faults are detected below the salt structure. We have described a mechanism where supra-salt faulting caused extension in the overburden and triggered reactive diapirism, breakthrough and extrusion of salt at the beginning of the Middle Keuper. Because overburden faulting is predominantly delocalised due to the decoupling of the salt layer, it can be focussed in weakness zones and therefore would prefer fractured crestal zones of salt pillows, if existent. Because intensified North Sea rifting in the Middle Keuper brought about major extension and faulting in the Ems region and the Glückstadt Graben (Ziegler 1990; Frisch and Kockel 1999), this concept of reactive diapirism is proposed to explain the breakthrough of mostly Triassic Glückstadt Graben and Ems region salt diapirs as well as for Jurassic to Early Cretaceous diapirism during the WNW-ESE oriented basin evolution. At those times when a syndepositional fault intersected and cut a near-emergent diapir crest, the caprock was breached and a pathway opened for salt extrusion coupled to renewed diapirism. Fault escarpments and faultinduced depressions allowed gravity-driven downward flow of salt across the landsurface to form a salt glacier (or namakier) and its subsequent preservation atop continental (arid-redbed) sediments that were accumulating in fault-defined collapse depressions adjacent to the growing salt diapir (Mohr et al. 2007). Somewhat similar fault intersections of overburden atop salt-filled bulges also control the onset of the namakier stage and active diapirism on or near the crests of compressional growth anticlines in SE Iran (Warren 2006). Additionally, bedded primary evaporites (halite) in near-diapir strata of peripheral sinks show extreme thicknesses of salt-sediment interbeds as can be observed in the

Middle Keuper sequences. We suggest that saline brine was supplied to a formerly arid landscape through diapir dissolution by groundwater. Diapiric growth and the developing relief of the peripheral sink controlled preservation and salt accumulation of the bedded salt units. The peripheral sinks of the diapirs usually preserve these characteristics, but below salt-overhangs they are difficult to detect on seismic. That the Late Triassic climate of the Northwest German Basin allowed terrestrial namakiers to be preserved and stacked with thicknesses that are seismically resolvable argues that hyperaridity dominated in these isolated extensional depressions across substantial time frames. If this were not the case, rainfall groundwater flushing would have prevented the preservation and stacking of namakier levels in the stratigraphy, as it possibly does in Iran and the Dead Sea depression today. This distinction supports the notion that hyperarid conditions were typical in the interior of the Mesozoic Pangaean supercontinent, even as it was rifting to become the incipient Atlantic Ocean. We expect that most of the diapirs in the CEBS have a long-term downbuilding history subsequent to diapiric breakthrough, which resulted in a complex salt-sediment interface and diapir geometry. Whilst older namakiers also occur in extensional terrestrial settings, today’s salt allochthons in extensional settings are largely restricted to submarine continental slope and rise settings in passive continental margins. Into-the-basin gravity gliding at the submarine continental margin, with upslope extension and downslope compression, controls the formation of diapirs and their allochthon wings and tiers. Contemporary salt crests and canopies typically lie beneath some 50-100 m of poorly consolidated deep-marine mud and silt, as in the Gulf of Mexico and the Santos Basin of Brazil (Warren 2006). Following reactive diapirism and diapiric breakthrough, we observed downbuilding, that is, the passive growth of a diapir near the surface, until Upper Keuper. Downbuilding usually dominates the evolution of most salt diapirs as a self-organised process, without the need of external forces, even if compression and extension play a modifying role (Rowan et al. 2003). Typical features are neardiapir sedimentary wedges, salt extrusions, unconformities and onlap-structures, salt re-sedimentation, increased thickness, unconformities and salt jags at the edge of the down-built sedimentary wedges. The salt jags we interpreted using seismics are sedimentary features of an extrusive diapir evolving from low accumulation rates in relation to the salt rise rates. This feature should not be confused with the phenomenon of wedge-shaped intrusions of Zechstein salt into Mesozoic salt layers caused by compressional tectonics (Kockel 1998), which is well

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known from parts of the basin which suffered strong inversion tectonics (Baldschuhn et al. 2001). Middle Jurassic to Early Cretaceous uplift and tilting, as a consequence of regional thermal doming and strikeslip movements (Ziegler 1990), is only indicated in the data by a major erosional unconformity (Base Cretaceous unconformity). Structures indicating local faulting or salt movements are not preserved. This is interesting, because an active, transpressive stress regime can be expected to have reactivated the growth of the covered salt structures. Modelling results are also consistent with a small influence of Late Cretaceous to Palaeogene compressional tectonics on salt movement in the case study area. The diapirs were shortened and uplifted whilst the overlying sediments were faulted and folded. Reactivation of salt movement is a possible response to thin-skinned compression and correlates well with the regional phases of

inversion in the German and Dutch basins (Kockel 2003; de Jager 2003). Decoupled compression responsible for Late Cretaceous pillow growth in the NE German basin (Kossow and Krawczyk 2002; Scheck et al. 2003a) leads to the squeezing of diapirs with narrowing of their necks, up-doming roofs and typical mushroom geometries in areas hardly affected by basin inversion. As a basinwide feature, late salt movement produced a positive relief of the base Late Cretaceous or base Tertiary above salt structures in relation to the regional datum. The approach of linking seismic techniques with structural restoration techniques and sedimentary sequence analysis was used in this chapter to unravel a complex, multiphase salt tectonic evolution in the course of changing stress fields in a defined area of the western CEBS. Each phase must have been a response to regional basin tectonics, which is the topic of the next chapter.

5

Chapter 5.4

Dynamics of salt basins M. Scheck-Wenderoth · Y. Maystrenko · C. Hübscher · M. Hansen · S. Mazur

5.4.1 Introduction Salt basins can occur in almost all types of plate-tectonic settings including intracontinental and cratonic basins, synrift basins, postrift passive margins as well as continental collision zones and foreland basins (Hudec and Jackson 2007; Warren this volume). Depending on the general setting, the dominant triggering factors for postdepositional salt mobilisation may be different. These factors may be gravitational loading, displacement loading and thermal loading (Hudec and Jackson 2007). Gravitational loading may result from laterally varying overburden thickness produced by sedimentation, de-

formation or erosion. Displacement loading results from the forced displacement of the boundaries of a salt body for example due to extension or compression and is especially important in the deformation of pre-existing salt structures. Thermal loading results from volume changes caused by changes in temperature. To complicate matters, the setting of a particular salt basin and the stress field governing the basin may change during its history. As detailed in the previous chapters, salt is, due to its mechanical properties, a very sensitive recorder of any deformation process affecting the basin. In other words, the history of salt tectonics can be of great help in understanding regional deformation.

Figure 5.4.1. Outlines of the area in which the Zechstein salt has initially been deposited in the Late Permian. The thickness and facies distribution of the underlying Rotliegend and the overlying Early Triassic sediments indicate that two individual sub-basins, the so-called Northern and Southern Permian Basins (Ziegler 1990), were separated by the Mid-North Sea – Ringkøbing-Fyn chain of structural highs during the evaporite deposition. Post-depositional mobilisation of the Zechstein salt from Mid Triassic times onward resulted in the complex thickness distribution of the salt at present

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As salt deforms far easier than clastic and carbonate sediments (Urai et al. this volume; Kukla et al. this volume), it acts as a decoupling agent between the cover layers and the basement during deformation. Thin-skinned deformation is the result of this mechanical decoupling of basement from cover in those areas where the original salt layer was thick with different wavelength and intensity of deformation above and below the salt. In contrast, thickskinned deformation affecting the entire column of the basin fill and the basement predominates where the salt layer thins towards and at the margins of the basin. The Central European Basin system (CEBS) is one of the best places in the world to study salt tectonics. Not only does it offer the full range of salt structures formed during a multiphase tectonic history, but it is also the place where the concepts of halokinetic deformation were first described. Since the early publications on salt tectonics in NW Germany (Trusheim 1960; Sannemann 1963, 1968) a large number of descriptions of salt structures has emerged and regional data compilations for the entire basin system facilitated a basin-scale understanding of processes related to salt movements. Consequently, the regional analysis of salt structures can unravel basin-scale controlling factors for phases of deformation in general. This chapter illustrates how such an analysis can help to understand the relationship between regional tectonics and salt movements. It provides an overview of the phases of salt mobilisation in response to tectonic triggering factors that affected the CEBS after the Zechstein salt was deposited. After a regional overview on the regional pattern of salt structures, detailed examples of key areas for each phase are illustrated. Finally, a brief summary of features that are characteristic to salt basins in an intracontinental setting attempts to highlight those findings that are transferable to intracontinental salt basins in general.

5.4.2 Regional pattern of salt structures in the CEBS The area where salt is observed today roughly corresponds to the area in which the Zechstein salt was initially deposited in the Late Permian (Fig. 5.4.1). The distribution of marginal carbonates as well as the thickness and facies distribution of the underlying Rotliegend and the overlying Early Triassic sediments (Bachmann et al. this volume ) indicate that the palaeogeoagraphic configuration of the Zechstein basin consisted of two individual sub-basins with WNW-ESE orientated axes. These so-called Northern and Southern Permian Basins (Ziegler 1990) were separated by the Mid-North Sea-Ringkøbing-Fyn chain of structural

highs during evaporite deposition and the Mid-North Sea-Ringkøbing-Fyn-High persisted as a structural high during most of the Mesozoic (Clausen and Pedersen 1999; Vejbaek 1997). The initial horizontal configuration of the salt, however, is only preserved in very restricted areas as post-depositional mobilisation of the Zechstein salt from Mid Triassic times onward resulted in the complex thickness distribution of the salt observed today. Looking at the presentday distribution of mobilised Zechstein salt in the CEBS (Fig. 5.4.2) it is obvious that the axes of salt structures largely follow the main structural grain. As detailed in chapter 3.3 (Kley et al. this volume), the structural pattern is dominated by two major directions, NW-SE and N-S, which are also present in the pattern of salt structures. Accordingly, salt structures with NW-SE directed axes are present parallel to the Teysseire-Tornquist Zone (TTZ) along the inverted Mid Polish Swell, to the SorgenfreiTornquist Zone (STZ) along the inverted northern margin of the CEBS and to the Elbe Fault System (EFS) along the inverted southern margin of the CEBS. These NW-SE trending salt structures developed above basement faults that were initiated prior to salt deposition. However, these structural elements were reactivated during Late JurassicEarly Cretaceous basin differentiation and played a key role during Late Cretaceous-Early Tertiary basin inversion. The second family of salt structures, characterised by roughly N-S orientated axes, is confined to the central parts of the CEBS and these salt structures align parallel to the margins of large grabens as the Glückstadt Graben, Horn Graben and Central Graben. The latter formed during the Mid-Triassic to Early Jurassic differentiation of the previous, WNW-ESE-orientated Permian basins into a series of sub-basins. Though almost the entire spectrum of salt structure types is found, areas of different deformation intensity can be distinguished. Most of the salt structures in the NE German sub-basin, between the large grabens in the Southern North Sea and the NW German sub-basin as well as in the western Norwegian-Danish sub-basin are developed as salt pillows. In contrast, salt diapirs and salt walls piercing their cover layers are present along the large N-S grabens, along the strongly inverted southern margin of the CEBS and the likewise inverted Mid Polish Swell as well as in the eastern Norwegian-Danish sub-basin, where the two structural trends intersect and interfere. Interpretation of seismic data revealed that the two families of salt structures evolved at different times and their maturity in terms of pillow-stage versus diapiric stage is strongly related to the amount of deformation affecting the sub-salt basement.

Dynamics of salt basins

Figure 5.4.2. Present distribution of salt structures in the CEBS and superimposed main faults (Maystrenko et al. 2006; Baldschuhn et al. 2001; Dadlez et al. 1998a; Evans et al. 2003; Jaritz 1987; Lokhorst et al. 1998; Nalpas and Brun 1993; Remmelts 1995; Scheck et al. 2003b). The mobilisation of the Zechstein salt resulted in two types of structures: (1) a province of NNW-SSE to NNE-SSW striking salt walls in the northern half of the Southern North Sea and the North German Basin initiated in Mid- to Late Triassic during accelerated subsidence in the large N-S trending grabens and (2) a province of NW-SE- striking salt structures in the southern half of the Southern North Sea and the North German Basin developed in Cretaceous times, coeval with the formation and inversion of NW-SE striking sub-basins. In the Polish basin the axes salt structures follow the NW-SE direction of the Teisseyre Tornquist Zone and have varying mobilisation ages. Profile numbers refer to the seismic sections shown in Figs. 5.4.3 to 5.4.5. CG: Central Graben, HG: Horn Graben, GG: Glückstadt Graben, RT: Rheinsberg Trough, AL: Aller Lineament, GE: Gardelegen Escarpment, BDF: Bornholm-Darlowo Fault Zone, KCF: Koszalin-Chojnice Fault Zone, PS: Pomeranian Segment of Mid-Polish Trough, KS: Kuiavian Segment, STZ: Sorgenfrei-Tornquist Zone, TTZ: Teisseyre-Tornquist Zone; EFS: Elbe Fault System, HCM: Holy Cross Mountains

5.4.3 History of salt movements in the CEBS Reconstructions of the initial salt thickness distribution for different sub-areas of the North German Basin suggest that thermal subsidence was the main controlling factor for basin evolution during the deposition of the Zechstein salt in Late Permian times (Maystrenko et al. 2005a,b; Scheck et al. 2003a,b). This regime persisted during the Early Triassic when the depositional area widened. Seismic data indicate continuous subsidence during the Early Triassic over large parts of the North German Basin (Hoffmann and Stiewe 1994; Kossow and Krawczyk 2002; Maystrenko 2005b; Maystrenko et al. 2006; Scheck et al. 2003a), of the Norwegian-Dan-

ish Basin (Clausen and Korstgard 1996; Clausen and Pedersen 1999; Vejbaek 1997), and of the Polish Basin (Dadlez et al. 1998b; Krzywiec 2002). This is expressed in continuous, parallel seismic reflections correlated with the Early Triassic Buntsandstein clastics and Muschelkalk carbonates that are generally not affected by syn-depositional faulting and continuously thicken towards the centres of Northern and Southern Permian Basins. In consequence, the Zechstein salt was covered by the Buntsandstein reaching up to several hundred metres in the Norwegian-Danish and up to a few thousand meters in the central part of the North German Basin (Baldschuhn et al. 2001; Maystrenko 2005b). The latter are overlain by some hundred meters of Muschelkalk carbonates characterised by a rather uniform thickness distribution.

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5.4.3.1 Salt movements in relation to Mid-Late Triassic regional extension

served in seismic data along northerly trending faults. This resulted in the formation of a series of large grabens and smaller, N-S orientated troughs, that contain the record of localised accelerated Mid-Triassic (Keuper) to Early Jurassic subsidence. Accordingly, locally increased thickness of Triassic deposits is found in these N-S trending depocentres (Fig.5.4.3). The most prominent examples are the Glückstadt Graben (Fig. 5.4.4a; Maystrenko et al. 2006; Baldschuhn et al. 2001; Kockel 2002), the Central Graben (Michelsen et al. 1987; Sundsbo and Megson 1993), the Horn Graben (Best and Kockel 1983; Clausen and Korstgard 1996; Clausen and Pedersen 1999), but also smaller depocentres as the Ems Graben (Mohr et al. 2005) or the Rheinsberg Trough (Scheck et al. 2003a). Thinning of the cover sediments in response to extension resulted in differential gravitational unloading of the salt layer above the thinned areas. At the same time accelerated subsidence and accompanying deposition in the forming grabens caused differential loading of the salt layer. Together, these processes initiated instabilities in the salt layer provoking salt rise parallel to the axis of the forming grabens and the formation of reactive diapirs (Vendeville and Jackson 1992a,b).

E-W-directed extension is deduced from the syn-depositional normal faulting in Triassic to Jurassic times ob-

Maximum extension is interpreted for Mid-Late Triassic Keuper times as recorded by a marked thickening of the

Though indications for tectonic activity (Beutler and Schüler 1987; Ziegler 1990) in the Early Triassic are sparse, traces of an incipient, E-W-directed extension are inferred from seismic data in the N-S-trending Glückstadt Graben, Ems Graben, in the Horn Graben and also in the graben system of the Central North Sea (Maystrenko 2005b; Mohr et al. 2005). This may have inititated first instabilities in the salt layer and thereby started the history of salt rise in the CEBS. Likewise, in the Mid-Polish Trough, some minor tectonic activity during Early to Mid Triassic times may have triggered initial salt movements (Dadlez 2003; Krzywiec 2002, 2006a,b; Hakenberg and Swidrowska 1997). These first indications of a new tectonic regime superposed on the declining thermal subsidence announce the subsequent, roughly E-W directed extension which controlled the structural evolution of the CEBS from Mid-Triassic to Early Jurassic times.

Figure 5.4.3. Thickness distribution of preserved Triassic deposits in the CEBS (compiled after Ziegler 1990; Vejbæk 1990; Baldschuhn et al. 2001; Scheck et al. 2003a,b; Bayer et al. 2002, Maystrenko et al. 2006). Note the locally increased thickness in N-S- orientated grabens and troughs. Only in the Polish Basin the NW-SE trend persisted

Dynamics of salt basins

Keuper deposits in the N-S-striking grabens and troughs. Localised accelerated Mid-Triassic to Early Jurassic subsidence was accompanied by salt mobilisation. Accordingly, most of the N-S-striking salt walls in the northern half of the Southern North Sea and the North German Basin were initiated in Mid- to Late Triassic times. In the Glückstadt Graben, for example, local stratigraphic thickening in the Keuper and Jurassic interval is visible in seismic

profiles across the graben (Fig. 5.4.4a) and some of these profiles also show evidence of syntectonic salt movement (Maystrenko et al. 2005a,b). The evolving depocentre was flanked by diapiric salt walls which were initiated above normal faults in the sub-salt basement. Salt rise led to the initiation of up to 200 km long salt walls, several kms high and several kms wide, parallel to the graben axis. The rising salt partially reached the palaeo-surface and was subjected

Figure 5.4.4. Seismic examples across the N-S striking structural elements active as Triassic-Jurassic extensional structures, affected by regional uplift in Mid-Late Jurassic to Early Cretaceous, by tectonic quiescence in the Late Cretaceous and by renewed subsidence in the Cenozoic. (a) Interpretation of migrated seismic section across the central Glückstadt Graben (profile 1 in Fig. 5.4.2, modified after Maystrenko et al. 2005a,b). Major salt movements occurred during the Keuper, when the Glückstadt Graben was affected by extension. Salt rise went so far that it extruded onto the palaeo-surface and was dissolved and re-deposited within the Keuper strata. Parts of the Jurassic were eroded in Late JurassicEarly Cretaceous times as indicated by the respective unconformity. The Late Cretaceous strata have an almost constant thickness and their parallel reflections pattern indicates a quiet tectonic setting with very minor salt movements in the Late Cretaceous. Salt movements resumed only in the Cenozoic. Stratigraphic key: P1-C-D = Undivided Lower Rotliegend, Carboniferous and Devonian deposits; P1(s) = salt-rich Rotliegend; P2 = Zechstein; P2+P1(s) = upper part of the Early Permian and Late Permian (undivided Zechstein and salt-rich Rotliegend); T1-2 = Lower Triassic and lowermost part of Middle Triassic (Buntsandstein); T2 = Middle Triassic without uppermost and lowermost parts (Muschelkalk); T2-3 = uppermost part of Middle Triassic and Late Triassic (Keuper); J = Jurassic; K = Cretaceous; Q-Pg = Palaeogene-Quaternary. (b) Interpretation of migrated seismic section across the western shoulder of the Rheinsberg Trough ending to the east in the trough centre (profile 2 in Fig. 5.4.2, modified after Scheck et al. 2003a). The onset of salt mobilisation is observed to be synchronous with the development of the NNE-SSW striking Rheinsberg Trough in the Late Triassic Keuper. This is inferred from stratigraphic thickening in the reflections interpreted as Late-Triassic (Keuper)-Jurassic towards the trough centre. The salt is almost completely removed below the trough. While normal faults are present in the Mesozoic cover, the base Zechstein appears as a strong, continuous signal below the trough and demonstrates that the evolution of the salt cover was independent from basement deformation. The Early Cretaceous rests unconformably on older units of conformal geometry and is truncated upward by an erosional unconformity. The latter separates Cenozoic from older units and indicates severe pre-Cenozoic uplift of the area

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Figure 5.4.5. (a) Interpreted seismic section and (b) detail of migrated and coherency filtered seismic line, running from the central North German Basin to the southern margin across the inverted Lower Saxony Basin (profile 3 in Fig. 5.4.2, modified after Mazur et al. 2005). This line illustrates continuous subsidence in the Early to Mid Triassic (Buntsandstein and Muschelkalk) represented by parallel, continuous reflections. The overlying reflections interpreted as Late Jurassic-Early Cretaceous show local stratigraphic thickening and indicate accelerated subsidence in the Lower Saxony Basin at the southern margin of the CEBS as compared to the Pompeckj Block further north. Late Cretaceous inversion resulted in a basement-involving uplift at the Lower Saxony Basin and folding of the salt cover in the NW German Basin. The tectonic decoupling by the Zechstein salt across large parts of the basin is obvious as well as localised, basement-involving deformation restricted to the Lower Saxony Basin, where salt is thin or absent

to leaching and re-deposition. A detailed case study of the Glückstadt Graben given later in this chapter illustrates the dynamics of these N-S structures. Salt withdrawal was a major factor for the creation of accommodation space not only in the Glückstadt Graben but also in most other N-Strending basins of the period (Baldschuhn et al. 2001; Best and Kockel 1983; Kockel 2002; Clausen and Pedersen 1999). Basement faulting is involved only in the largest of the N-S grabens, whereas in many areas the thinning of the salt cover was sufficient to initiate reactive diapirs. A good example of a reactive diapir with thin-skinned cover deformation is the Rheinsberg Trough in eastern Germany (Fig. 5.4.4b). Here, salt withdrawal almost completely balanced extension in the salt cover as indicated by seismic data and subsidence restoration (Scheck et al. 2003b). Though the structural pattern of the North German and the Norwegian-Danish basins illustrates that the tectonic regime changed in Triassic-Jurassic times to E-W-orientated extension, the structural trend is not observed to have changed in the Polish Basin. Moreover, in the Mid-Polish Trough the NW-SE-orientated axis of maximum subsidence and sedimentation persisted. This phenomenon nicely illustrates how a deep-seated zone of crustal weakness, in this case the TTZ located below the Mid-Polish Trough, can exert a dominant control on the deformation pattern (Dadlez et al. 1998a; Krzywiec 2006a,b). Yet, intense Late Triassic salt movements are also observed within the Mid-Polish Trough and salt structures with NW-SE-trending axes formed parallel to the TTZ. These salt movements resulted in the formation of salt rim synclines filled with Mid-Late Triassic deposits (Krzywiec 2006a,b; Mazur et al. 2005).

5.4.3.2 The salt during Jurassic-Early Cretaceous basin differentiation During the Mid Jurassic large parts of the central CEBS were affected by regional uplift which appears to have been diachronuous across the CEBS and the underlying mechanisms are still enigmatic (Underhill and Partington 1993; Vejbaek 1997; Baldschuhn et al. 2001; Jaritz 1987; Surlyk and Ineson 2003; Kossow and Krawczyk 2002; Maystrenko et al. 2005a; Scheck-Wenderoth et al. in press). The effect on salt stability, however, was most probably that of a regional gravitational unloading of the northern half of the North German (Pompeckj Block; Jaritz 1987) and Southern North Sea basins. It is very difficult to asses the detailed deformation pattern resulting from this event as syn-kinematic sediments are either lacking or have been eroded. Remnants of preserved Jurassic salt rim synclines imaged in seismic data from the central part of the CEBS indicate that salt rise has continued throughout the Jurassic parallel to the N-S-trending grabens and troughs. Successive depletion of the salt layer beneath the axial parts of these grabens and troughs caused a wavelike outward propagation of rising salt walls as illustrated in the case study of the Glückstadt Graben (Fig. 5.4.4a). This resulted in increasingly younger rim synclines outward and most of the preserved Jurassic rim synclines are located marginal to the ones that had formed in the Keuper. Amazingly, the areas marginal to the central and uplifted part of the CEBS were continuously subsiding throughout the Triassic and Jurassic. Moreover, these marginal areas were even affected by accelerated subsidence. Accordingly, NW-SE-orientated sub-basins developed along the

Dynamics of salt basins

Figure 5.4.6. (a) Interpreted seismic section DEKORP BASIN9601 (profile 4 in Fig. 5.4.2), crossing the NE German Basin from south of Rügen to the southern basin margin and (b) detail of migrated seismic line. The line is perpendicular to the Permo-Triassic basin axis as well as to the WNW-ESE striking inversion structures, but parallel to the NNE-SSW striking Rheinsberg Trough and indicates continuous subsidence during Triassic-Jurassic. At the southern margin localised subsidence is indicated for the Early Cretaceous and inversion for the Late Cretaceous. Deformation is tectonically decoupled by the Zechstein salt below the basin whereas basement deformation is localised and most intense at the Elbe Fault System. There, the base Zechstein is displaced by 2 seconds Two Way Traveltime (s TWT) along the Gardelegen Escarpment (GE), a sub-fault of the Elbe Fault System. Some smaller basement faults are visible to about 50 km north of the Gardelegen Escarpment and at the northern margin beneath the Grimmen High. Over the largest part of the basin the base Zechstein appears as a strong, continuous signal below salt-depleted areas as well as below salt pillows and diapirs

marginal areas of the basin system in the Late Jurassic to Early Cretaceous, whereas the central part of the CEBS remained in an uplifted position. This local subsidence was enhanced by coeval normal faulting and coeval salt movements. Thus a second family of salt structures with NW-SE orientated axes (Fig. 5.4.2) was initiated in the latest Jurassic to Early Cretaceous times. The Lower Saxony Basin is a typical example for this type of structure. A seismic section, perpendicular to the basin axis (Fig. 5.4.5), illustrates this accelerated subsidence indicated by local stratigraphic thickening of the reflection packages interpreted as Jurassic and Early Cretaceous. Salt structures parallel to the Aller Lineament show pronounced stratigraphic thinning of the Jurassic to Early Cretaceous interval above their crests and are cut at their northern flank by steep faults. At the southern end of the profile, where no salt is present, normal faults offsetting the entire sediment column are observed and indicate that a regional change of the tectonic stress field (with respect to the foregoing E-W extension) played its part in the new deformation regime and related salt mobilisation. Similar elongated, NW-SE-trending depocentres accompanied by parallel salt structures are observed in other areas along the southern margin of the CEBS including the Sole Pit Basin, Broad Fourteens Basin, the West and Central Netherlands Basins, the Altmark Basin and the Subhercynian Basin (Betz et al. 1987; Jaritz 1980, 1987; Nalpas et al. 1995; Scheck et al. 2002a). Analogously, such structures are reported parallel to the STZ at the northern margin of the Norwegian-Danish Basin (Surlyk and Ineson 2003; Vejbaek 1997) and from the Polish Basin (Dadlez et al. 1995, 1998a; Krzywiec 2002, 2006a,b).

5.4.3.3 Salt movements during Late Cretaceous-Early Tertiary compression After experiencing localised subsidence in the Early Cretaceous, the NW-SE-orientated sub-basins along the margins of the CEBS as well as the axial part of the Polish Basin were the sites of localised uplift during the latest Cretaceous. Coeval with the inversion of the NWSE-striking sub-basins, a new phase of salt movement is observed which enhanced the growth of NW-SE-striking salt structures. Especially in the NE German Basin, it was during the Late Cretaceous – Early Tertiary that the salt structures reached the piercing state, mostly as a result of compression. Likewise, local uplift accompanied by salt rise took place at the northern margin of the CEBS along the STZ (Vejbaek 1997) and along the inverted Mid Polish Swell parallel to the TTZ (Krywiecz 2006b). Seismic data perpendicular to the EFS at the southern margin of the CEBS (Figs. 5.4.5 and 5.4.6) indicate that many of the reverse or transpressive faults accommodating uplift show normal or transtensional activity during the deposition of the underlying Late Jurassic-Early Cretaceous sediments and that the area with stratigraphic thickening of the Early Cretaceous coincides with the area of maximum uplift during the Late Cretaceous. These phenomena are also reported from the Danish Central Graben (Cartwright 1989; Vejbaek and Andersen 2002), the Sole Pit Basin (Badley et al. 1993; Buchanan et al. 1996; Nalpas et al. 1995), the Broad Fourteens Basin (De Lugt et al. 2003; Nalpas et al. 1995), and the Subher-

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M. Scheck-Wenderoth · Y. Maystrenko · C. Hübscher · M. Hansen · S. Mazur Figure 5.4.7. (a) Simplified tectonic map of the Glückstadt Graben with location of the seismic line shown in Fig. 5.4.4a (position of salt domes after Baldschuhn et al. 2001). (b) 3D structural model of the Glückstadt Graben (modified after Maystrenko et al. 2006) and adjacent areas (for location see blue frame in a)

Dynamics of salt basins

cynian Basin (Franzke et al. 2004; Kossow and Krawczyk 2002; Otto 2003; Voigt et al. 2004). Basement-involving compressive deformation is localised along the margins of the CEBS where Zechstein salt is thin or missing. In contrast, compressive deformation in the central part of the CEBS is mostly concentrated above the Zechstein salt. This deformation is expressed in folding of the salt cover with salt-cored anticlines and synclines from where the salt has been withdrawn. The intensity of basement-involving deformation decreases rapidly with increasing distance from the marginal fault systems and the base Zechstein is almost continuous below large parts of the basin area. Similar inversion structures have been described along the STZ at the northern margin of the Norwegian-Danish Basin (Vejbaek and Andersen 2002; Vejbaek 1997) and along the Mid Polish Swell. At the Mid Polish Swell deformation is decoupled by the salt layer in the Kuiavian and Pomeranian segments whereas it is thick-skinned in the Holy Cross Mountains in absence of salt (Erlström et al. 1997; Krzywiec 2002, 2006a,b; Mazur et al. 2005). In contrast to the NW-SE trending structures, most of the older (Triassic), N-S orientated structures were not inverted and their parallel salt structures experienced a phase of stability during the Cretaceous. This is expressed in constant Cretaceous thickness over the crests of these salt structures (Figs. 5.4.4).

deepest structures of the CEBS and contains up to 12,000 m of Permian, Mesozoic and Cenozoic sediments. Structurally, the Glückstadt Graben can be subdivided into three domains (Fig. 5.4.7): (i) the Central Glückstadt Graben, (ii) the marginal Westholstein, Eastholstein and Hamburg troughs that are of Jurassic to Cenozoic age — and (iii) the flanks of the basin — the Westschleswig and EastholsteinMecklenburg blocks. Its structure is strongly complicated by NNE-SSW-striking, elongated salt walls which formed during post-Permian times. Therefore, one of the main deformation mechanisms in the area of the Glückstadt Graben is salt tectonics. Details of the evolution are expressed in a variety of salt structures representing different stages of growth such as salt rollers, anticlines, pillows, stocks, and, most pronounced, elongated salt walls.

The Cenozoic deposits are separated from the Mesozoic by an unconformity, which is erosional in character over large parts of the CEBS. However, after this erosional event, sedimentation resumed in the CEBS and again, the orientation of depositional axes changed with respect to the foregoing tectonic phase. In contrast to the NW-SE trend of the Late Cretaceous phase of inversion, the dominant structural trend of depocentres in the Cenozoic is N-S orientated and thus similar to the earlier phase of Triassic extension. Moreover, newly-formed N-S orientated normal faults, as for example observed in the Cenozoic of the Glückstadt Graben (Figs. 5.4.4), additionally prove that E-W directed extension controlled this youngest tectonic phase of the CEBS. Older, N-S orientated salt walls, which had remained quiescent during the Late Cretaceous compressive phase were reactivated and experienced a new phase of growth (Figs. 5.4.4).

At the end of the Early Triassic (late Buntsandstein), incipient E-W directed extension caused the formation of a very narrow trough within the axial part of the Glückstadt Graben (Brink et al. 1992; Kockel 2002). However, the main phase of extension is recorded in the Late Triassic (Keuper) sediments, deposited in a broad and deep trough bounded by rising salt walls (Kockel 2002; Maystrenko et al. 2006). During the Jurassic, three NE-SWaligned troughs (Eastholstein, Westholstein and Hamburg; Fig. 5.4.7) developed marginal to the Central Glückstadt Graben indicating a predominantly extensional regime in the area. The post-Jurassic sequence is bounded by an Early Cretaceous unconformity (Jaritz 1969). During the Late Cretaceous, the sediments were deposited in a gentle platform-type depression that extended far beyond the boundaries of the Glückstadt Graben. The Late Cretaceous-Early Cenozoic phase of basin inversion observed in other parts of the CEBS is not documented in the Glückstadt Graben. In contrast to the NW-striking reverse faults and flexures that developed along the southern and northern margins of the CEBS or along the TTZ in the Polish Basin (Ziegler 1990; Scheck et al. 2002a; Otto 2003; Scheck-Wenderoth and Lamarche 2005; Mazur et al. 2005; Vejbaek 1997), the Glückstadt Graben appears to have been tectonically quiet and was covered by an Late Cretaceous layer of uniform thickness. Only at the marginal parts of the Glückstadt Graben indications for increased syn-inversion salt activity during Cretaceous-Early Tertiary compression are observed (Kockel 2002; Maystrenko et al. 2006). In the Cenozoic, normal faulting and rapid subsidence in the Eastholstein, Westholstein and Hamburg Troughs indicate renewed tectonic activity (Maystrenko et al. 2005a,b).

5.4.4 Case Study Glückstadt Graben

5.4.4.1 Structural features of the Glückstadt Graben

The Glückstadt Graben is a key area to illustrate the interplay of extension and reactive diapirism. It is one of the

The structural features of the Glückstadt Graben are best described along a regional reflection seismic profile which

5.4.3.4 Salt movements in relation to Late Tertiary regional extension

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east, and is covered by post-Permian sediments of more than 2 s TWT thickness. This structural zone has small dip gradients (sometimes almost horizontal) at the base of Zechstein, and the salt and its overburden appear relatively undeformed.

Figure 5.4.8. Map summarising changing subsidence with time within the Glückstadt Graben. Centres of maximum sediment thickness indicate a shift of deposition from the central part in the Triassic to the marginal parts in the Quaternary-Neogene (modified after Maystrenko et al. 2005a, 2006)

crosses the central and eastern parts of the area of consideration. (Figs. 5.4.4a). The thickness differences shown in the profile distinguish between the two main structural zones: the Triassic Central Glückstadt Graben with the superimposed Jurassic and Cenozoic Eastholstein Trough and the Eastholstein-Mecklenburg block. In the Central Glückstadt Graben, the base of Triassic sediments is located at more than 4.5 s TWT (Fig. 5.4.4). The uppermost Middle-Late Triassic (Keuper) represents the largest part of the infill of the Glückstadt Graben and shows increasing thickness towards the basin centre. There, the reflection patterns show evidence of strong syn-depositional thickening of the Keuper (up to 2.5 s two way traveltime ≈ 5800 m) compared with the underlying strata. The presence of baselapping strata of the thick Keuper sequence indicates rapid subsidence, accompanied by downbuilding. At its south-eastern margin, the Triassic depocentre is separated from the Eastholstein-Mecklenburg block by the Eastholstein trough, which displays increased thicknesses of Jurassic and Cenozoic sediments. The Cretaceous sequence shows an almost constant thickness along the line with a smooth thinning from the Eastholstein Through towards the SE flank. The Eastholstein-Mecklenburg block bounds the Glückstadt Graben to the south-

Therefore, this seismic line illustrates that the axial parts of the thickened sediments at the different stratigraphic levels are rarely vertically aligned (see white arrows in figure 5.4.4). This indicates that the location of depocentres shifted in time and in space. For instance, thickened Keuper is bounded by a thick Jurassic sequence in the western part of the Eastholstein Trough. This Jurassic depression is characterised by thickening of strata, indicating rapid subsidence in comparison to other parts of the section. The thickened Jurassic is covered by thick Cenozoic but without vertical alignment of the axes of maximum thickness. This indicates that the salt cover was deformed due to salt movements and that the salt decoupled deformation of the overburden from the strata below. Furthermore, deformations of the overburden are observed above the approximately flat base salt over the eastern part of the line. This again shows that the salt cover was deformed due to salt movements and that the observed thickening of the sediments must have been associated with simultaneous salt movements. Across the whole basin, the present-day distributions of sediments at different stratigraphic levels show that the centre of sedimentation moved away from the central part of the original Triassic trough towards its margins (Fig. 5.4.8). It is obvious that the marginal troughs (Westholstein, Eastholstein and Hamburg) are separated by thick Triassic (mainly Keuper) deposits which are strongly thickened towards the axial part of the Glückstadt Graben (Fig. 5.4.4). On the other hand, the area of the Central Glückstadt Graben is characterised by relatively thin Cretaceous and Cenozoic sediments and partly by the absence of the Jurassic (Fig. 5.4.4). Furthermore, the distribution of thickness maxima demonstrates that the thick Jurassic is covered by thickened Cretaceous, Palaeogene and Neogene but without vertical alignment of those axial parts within the marginal troughs (Fig. 5.4.8). Therefore, figure 5.4.3 suggests that a greater amount of subsidence occurred close to the active salt structures, and may have resulted in a gradual depletion of Permian salt from the source layer.

5.4.4.2 3D reconstruction of salt movements Figure 5.4.9 illustrates the outcomes of the 3D inverse modelling of the spatial and temporal evolution of the salt structures in the Glückstadt Graben (Maystrenko et al. 2006). The theory of this approach (Scheck and Bayer

Dynamics of salt basins

Figure 5.4.9. Modelled thickness of the Permian salt in the Glückstadt Graben projected on 3D views of the top Permian salt for different stages of backstripping (after Maystrenko et al. 2006). Isopach maps show the evolution of salt structures from the end of the Triassic (a) to present-day (f). Note the gradual migration of the depletion zone from the central part of the Glückstadt Graben (a) towards its margins (f) shown by the light blue colour. CGG – Central Glückstadt Graben, EHT – Eastholstein Trough, HT – Hamburg Trough, WHT – Westholstein Trough

1999; Scheck et al. 2003b) considers salt flow as a consequence of spatially changing overburden load and requires several steps. The first step is to remove the sediments above the selected stratigraphic level. Then the remaining sediments are decompacted. Subsequently, redistribution of the salt is calculated for the reduced overburden load. The behaviour of salt is assumed to be similar to a viscous fluid and salt volume is conserved. After salt redistribution, a load balance at the base salt is achieved so far as salt and overburden are in hydrostatic equilibrium. Finally, the isostatic response of the entire model is calculated according to the new mass distribution using an Iceberg approach (Kley et al. this volume; Box 3.3.2.). The described sequence of modelling was applied to reconstruct the geometry of salt structures from present day back to the end of the Triassic. One of the most remarkable results is the expansion of the depletion zone shown by the light blue colour in figure 5.4.9. It can be seen that the Permian salt layer was almost depleted within the central part of the basin in the Triassic (Fig. 5.4.9a). A second feature is the increasing number of linear zones of increased salt thickness shown in yellow or different shades of red. These zones represent growing salt struc-

tures. Some of these salt structures were already in place in the Triassic Central Glückstadt Graben, whereas the marginal troughs (West-, Eastholstein and Hamburg) are characterised by locally increased salt thickness related to younger salt movements. During the Jurassic, depletion of the salt layer affected the SW and NE margins of the Central Glückstadt Graben (Fig. 5.4.9b). Little change is visible for the Early (Fig. 5.4.9c) and Late Cretaceous, only some reductions of salt thickness occurred within the Westholstein and Hamburg Troughs (Fig. 5.4.9d). In contrast, the thickness of the salt was strongly reduced within the West- and Eastholstein Troughs during the Palaeogene (Fig. 5.4.9e) and further expansion of the salt depletion zone towards the basin flanks took place during the Quaternary-Neogene (Fig. 5.4.9f). The results support the conclusion that initial salt movements took place during the Triassic within the central part of the Glückstadt Graben (Fig. 5.4.9a). Following this initiation, strong salt tectonics occurred at the margins of the former Triassic Graben in the Jurassic (Fig. 5.4.9b), while during the Cretaceous-Tertiary, additional growth of salt structures took place within the marginal troughs (Figs. 5.4.9c-f). Therefore, the 3D inverse modelling pro-

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vides evidence for a continued depletion of the salt layer progressing from the central part of the basin outwards to the basin flanks. The development of this depletion zone (Fig. 5.4.9) correlates perfectly with the observed distribution of the maximum thickness and therewith the development of subsidence centres for the different stratigraphic units (Fig. 5.4.8). In 2D, this phenomenon has been observed by Sannemann (1968) who introduced the term “salt-stock families” - salt diapirs of the Glückstadt Graben spread in time becoming younger by moving away from the axial part towards the basin flanks. Thus, the modelling results suggest that salt withdrawal during the Meso-Cenozoic may be considered as one of the main controlling factors for the subsidence in the Central Glückstadt Graben and the subsequent Jurassic-Cenozoic marginal troughs (Hamburg, East- and Westholstein). In other words, the progressive depletion of the salt layer, from the central part towards the margins, created much of the accommodation space for sedimentation along with tectonic subsidence in the Glückstadt Graben.

5.4.4.3 Salt movements in relation to tectonic events The good temporal correlation between the main phases of salt movements and tectonic events implies that salt mobilisation was tectonically induced in the Glückstadt Graben. The first stage of salt movements in the Central Glückstadt Graben (Figs. 5.4.4 and 5.4.9a) could have been triggered by Early Triassic extension that culminated during the Keuper (Brink et al. 1990, 1992; Kockel 2002; Maystrenko et. al. 2005a,b). This is consistent with the assumed regional stress field during the Triassic, characterised by E-W extension and responsible for normal faulting of the salt base with formation of N-S striking graben structures as described above. This regional Triassic extension may have triggered reactive diapirism and caused the formation of the deep primary rim synclines. Once the salt structures had reached a critical size, buoyancy forces (Schultz-Ela et al. 1993) supported their continued growth until the Jurassic when extension-induced regional stresses once more affected the Glückstadt Graben. The Jurassic pulse of salt activity (Figs. 5.4.9b), correlates temporally with extensional tectonics south of the Glückstadt Graben, where accelerated subsidence is observed within the Pompeckj Block and the Lower Saxony Basin (Kockel 2002). In most cases, Jurassic salt movements caused the formation of secondary rim synclines adjacent to the Keuper primary depocentres. This deformation persisted into the earliest Cretaceous (the Berriasian and the Valanginian) when salt continued to move close to the former Jurassic salt structures. During the post-Valanginian, almost all salt stocks and walls were covered by Early Cretaceous (Hauterivian-Albian) deposits, the almost constant thick-

ness of which indicate an absence of strong salt movements in the Glückstadt Graben. Similar remarks can be made for the lower part of the Late Cretaceous which is also characterised by approximately constant thickness. This suggests that salt movements ceased during the late Early Cretaceous-early Late Cretaceous possibly due to a combination of tectonic quiescence and partial depletion of the source salt layer. Renewed salt movements, especially within the marginal troughs, during the latest Late Cretaceous-Early Cenozoic correlate temporally with compressive deformation. In the Hamburg Trough, stratigraphic thickening of the uppermost Late Cretaceous-Early Palaeogene (Fig. 5.4.8) implies that the salt was active during this period, probably as a result of compressional stresses related to the Alpine Orogeny (Kockel 2002). Possibly the thickskinned deformations observed at the Aller Lineament in the latest Late Cretaceous-Early Cenozoic, (Mazur and Scheck-Wenderoth 2005) were transmitted northward into the overburden of the Glückstadt Graben causing salt movements without involving the salt base into the deformation. The youngest phase of Palaeogene-Neogene salt withdrawal led to the growth of almost N-S orientated salt structures mainly at the margins of the basin (Fig. 5.4.9e,f). This phase of salt tectonics correlates temporally with roughly W-E directed extension, indicating a renewed change in tectonic regime after Late CretaceousEarly Cenozoic compression.

5.4.5 Case Study NE German Basin Like in other parts of the CEBS, phases of salt tectonic deformation were controlled by the regional stress field in the NE German Basin (NEGB). However, this subbasin reacted differently in the sense that the salt basement remained more or less stable in the basin area and deformation of the salt cover is an outstanding example of thin-skinned tectonics. Basement-involving deformation is only observed beneath the inverted Grimmen High at the northern margin of the NEGB and beneath the severely uplifted Flechtingen High at the southern margin of the NEGB (Fig. 5.4.6). Both areas coincide with areas of long-lived crustal weakness. The Grimmen High extends along the north-western prolongation of the Mid Polish Swell and thus may be influenced by the TTZ. The Flechtingen High is bounded by sub-faults of the EFS, a regional fault system structuring the entire southern margin of the CEBS. Apart from the basin margins the deformation appears to be completely decoupled within the NE German Basin. The reflector interpreted as base Zechstein appears as a strong, continuous signal in seismic data below both salt-depleted areas as well as below salt pillows and diapirs

Dynamics of salt basins

Figure 5.4.10. Results of backstripping in the NE German Basin (Scheck et al. 2003b).The calculations are based on a structural model resolving Early Permian Rotliegend, Late Permian Zechstein salt, EarlyTriassic Buntsandstein and Muschelkalk, Late Triassic Keuper, Jurassic, Early Cretaceous, Late Cretaceous and Cenozoic. The backstripping results indicate that major changes in salt configuration took place during Late Triassic –Jurassic extension and during Late Cretaceous inversion with most intense deformation during Late Cretaceous inversion

(Figs. 5.4.4b; 5.4.6). In contrast, the cover sediments of the salt are deformed though the salt structures are generally less mature than in other parts of the CEBS. The first significant tectonic event in this sub-basin also was E-W-extension in Middle–Late Triassic times. In consequence, the NNE–SSW striking Rheinsberg Trough has developed in the Late Triassic perpendicular to regional extension (Scheck and Bayer 1999) and continued to subside during Early and Mid Jurassic times. Above an intact salt basement, the Rheinsberg Trough contains up to 1500 m of Keuper deposits and formed as a major salt rim syncline, 180 km long in the NNE-SSW direction, and 70 km wide, (Scheck et al. 2003a). Sediment downbuilding in the through centre (Fig. 5.4.4b) was accompanied by salt withdrawal from beneath this area and by salt rise marginal to the through. A second phase of deformation in the Late Jurassic and Early Cretaceous is recorded as a rotation of depocentral axes to NW-SE in the NE German Basin. The NW–SE striking Subhercynian and Altmark–Brandenburg Basins formed together with salt structures characterised by parallel axes (Schwab 1985; Scheck and Bayer 1999). Sub-

sequently, Late Cretaceous –Early Cenozoic compression led to uplift and erosion in the NEGB with differential inversion along NW–SE striking structures most intense along the northern and southern margins. At the basin margins compressive deformation is thick-skinned and expressed as steep reverse faults. In the basin area, once more only the salt cover is folded with salt-cored anticlines of decreasing amplitude and wavelength northward. The Late Cretaceous also was the time when the salt structures in the southern part of the NE German Basin reached the diapiric stage and pierced their cover layers. An erosional unconformity delimits the top of the Mesozoic sequence indicating relative uplift during or after the compression. Finally, renewed subsidence during the Cenozoic is documented in the sedimentary record of the NE German Basin, again accompanied by salt rise. The wavelength of Cenozoic salt rim synclines is larger than in the foregoing Cretaceous phase of compression and Cenozoic rim synclines partially cut and overstep the Cretaceous ones. The observed decoupled deformation suggests that salt movement and related formation of rim synclines may have balanced tectonic stresses in the salt cover. Results from 3D backstripping, considering salt flow as a

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Figure 5.4.11. Distribution of the calculated vertical tectonic movement for a structural model of the SW Baltic Sea area (Hansen et al. 2007) for the Middle Triassic (a), the Late Triassic and Jurassic (b), the Cretaceous (c) and the Cenozoic (d). Positive values indicate tectonic subsidence and negative values uplift. Present topography (e) and predicted future topography (f) resulting after isostatic balance of the present-day setting. Coordinates are UTM33N [m]

consequence of spatially changing overburden load distribution, isostatic rebound and sedimentary compaction for each backstripping step, (Fig. 5.4.10; Scheck et al. 2003a,b) support this hypothesis. Tectonic stresses related to regional E-W extension in the Triassic may have been transmitted in the weaker salt cover, but not in the more consolidated salt basement. Thus extensional unroofing may explain the development of the Rheinsberg Trough and the related reactive salt rise during this period. Likewise, Late Cretaceous-Early Cenozoic strain localisation along the SW margin of the NE German basin may have propagated into the salt cover of the basin and the resulting folding may have been balanced by viscous salt flow into the cores of anticlines.

5.4.6. Case Study SW Baltic Sea A detailed 3D structural model of the south-western Baltic Sea and the adjacent onshore areas (Hansen et al. 2005, 2007) provides the link between the thickskinned deformation of the Glückstadt Graben with its large salt walls and an area of very weak salt mobilisation east of the Glückstadt Graben in the NE German Basin. This model integrates six successions: Late Permian Zechstein; Early Triassic; Middle Triassic; Late Triassic–Jurassic; Cretaceous and Cenozoic. Like for the other two case studies this structural model was the basis for a 3D backstripping approach considering salt redistribution. With the approach the amount and re-

Dynamics of salt basins

gional distribution of tectonic subsidence or uplift can be determined. Two periods of accelerating salt movement in the area have been found. The first one has been correlated with the E–W directed extension during the Late Triassic – Early Jurassic. Accordingly, the tectonic subsidence distribution pattern shows that tectonic subsidence was small and rather uniform over most of the study area for the Late Triassic (Keuper) and the Jurassic (Fig. 5.4.11b), but strongly increases in the south-westerly direction towards the Glückstadt Graben. This pattern of localised subsidence in the Glückstadt Graben can also be traced throughout the younger stratigraphic levels and is always accompanied by parallel regions of uplift delineating rising salt walls. The second phase of salt movement correlates with the Late Cretaceous–Early Cenozoic inversion. Apart from the dominant signal of the subsiding Glückstadt Graben, the pattern of tectonic subsidence for the Cretaceous indicates a weak structuration with a NWSE trend and the tectonic subsidence determined for the Cenozoic (Fig. 5.4 11c) reveals uplift in the north-eastern part of the study area along a wider region trending NWSE. This region coincides with the Grimmen High, the Late Cretaceous-Early Cenozoic uplift of which is confirmed by seismic data (Fig. 5.4.6).

both the boundaries of these blocks as well as the contrasting properties of the blocks with respect to each other can determine the localisation of salt tectonic deformation. It is therefore crucial to understand the crustal structure beneath intra-continental basins to correctly assess domains of different salt tectonic structural provinces.

5.4.7 General findings for salt-containing intra-continental basins

Different parts of the CEBS show different intensities of salt-tectonic deformation in response to the changing regional stress conditions. As described above, some areas display a strong localisation of deformation, whereas also parts exist within the basin system that show far weaker deformation or even preserved the initial depositional geometry of the salt. Though there is a clear correlation between thick-skinned deformation and thin initial salt thickness as well as between thin-skinned deformation and large initial salt thickness, the lateral variation in overall deformation intensity does not always correlate with the initial thickness of the salt layer. In general, deformation intensity increases towards the basin margins in the North German and Norwegian-Danish basins as well as towards the axial part of the Polish Basin. This phenomenon is most likely related to the presence of zones of crustal weakness in theses areas. In particular, the STZ-TTZ and the EFS are structures that have been repeatedly the site of strain localisation since at least Late Carboniferous times. Accordingly, they were the places where strong deformation took place. Likewise, localised deformation involving the salt basement is observed along the N-S trending Glückstadt Graben, Central Graben and Horn Graben. The repeated activity of these N-S-structures indicates that they also represent zones of crustal weakness. As a second trend, deformation intensity decreases from west to east and fades out in the NE German Basin. There, the salt basement appears to be undisturbed over the largest part of the NE German Basin and the amplitude of salt structures decreases toward the Baltic Sea. Over large parts of the Baltic Sea and east of the Glückstadt Graben the salt is hardly mobilised at all though these areas are characterised by a large initial thickness of the salt. This observation confirms that buoyancy alone does not initiate salt rise and that other triggering mechanisms are required to mobilise the salt. In addition, the observed deformation pattern indicates that the salt basement is segmented into domains of different strength with weaker provinces prone to strain localisation and stronger provinces with a yield strength larger than the effective stress.

In addition to the physical laws that apply in all plate-tectonic settings, some controlling factors can be eliminated that are characteristically influencing the post-depositional mobilisation of salt in intra-continental basins. One of these factors is the segmentation of the salt basement into domains of different rheological properties. As intra-continental basins develop above a mosaic of different crustal blocks,

These findings seem to apply also in other intra-continental basins as for example in the Dnepr-Donetz Basin, where the deep crustal structural grain determines the location and orientation of salt structures, but also the distribution of deformation intensity (Stovba and Stephenson 2003). Furthermore, many salt structures display renewed growth after a regional interruption of salt movements

An interesting aspect of this work is the result from a neotectonic forward modelling in an attempt to predict the future topography when the system is in tectonic equilibrium. This was completed by letting the starting model run towards equilibrium in order to determine the neotectonic vertical movement in the south-western Baltic Sea. The resulting surface (Fig. 5.4.11f) shows where the coastline would be running when the system was in isostatic equilibrium. For comparison the present-day topography is presented in figure 5.4.11e. The result revealed that many of the salt structures in the region are still active. If the present-day load regime would continue towards equilibrium, the area around the western Baltic Sea would be more or less equally subdivided in one land and one sea area and the future coastline would run with a WNW–ESE trend.

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and seem to be controlled by similar mechanisms as in the CEBS. In summary it can be concluded that in the CEBS individual basement blocks react differently. The North Sea-NW German Basin area, for instance, is characterised by a regularly spaced distribution of salt structures and faults affecting the salt basement, though the wavelength of structures is different above and below the salt. In contrast, the NE German Basin reacted a stable block, that experienced regional subsidence and uplift and is affected by faults only at its margins. Salt tectonic deformation is far weaker and salt structures are less mature within this block than in the North Sea and in the NW German Basin. The Polish basin displays a third type of salt-tectonic deformation which is clearly dominated by the deep crustal boundary of the TTZ. The structural evolution of this salt tectonic province shows that if deep reaching crustal dis-

continuities are present they may control the orientation of forming salt structures to the same degree as the regional stress field. The proximity to this type of basement boundary also appears to control deformation intensity. Another interesting attribute of intra-continental salt basins is that they can store the effect of different stress fields. The orientations of salt structures initiated in different tectonic phases often are preserved although the causative stress field has long declined. Different generations of salt structures therefore are very good indicators for the regional deformation history. As sedimentation takes place in balance with salt withdrawal and water depth in intra-continental basins is generally very shallow, good reconstructions of the subsidence history can be obtained by backstripping. However, processes of salt extrusion and dissolution may be important and should be considered in these approaches.

5

Chapter 5.5

Temperature fields, petroleum maturation and fluid flow in the vicinity of salt domes F. Magri · R. Littke · S. Rodon · U. Bayer · J.L. Urai

5.5.1 Introduction Salt domes (diapirs) are common geological features of sedimentary basins. Owing to their physical and structural properties, salt diapirs strongly influence the temperature field of sedimentary basins. This has a great impact on the maturity of organic matter and the timing of hydrocarbon generation. Furthermore, temperature disturbances in the vicinity of salt domes strongly control groundwater transport processes since they provide the coupling between hydraulic and thermally-driven forces. This chapter consists of three parts. In the first part, the temperature anomalies around salt structures are illustrated. An example based on a salt diapir of the Central European Basin System (CEBS) will elucidate the effects of thermal anomalies on oil maturation. In the second part, an example of fluid flow inside salt structures is given,

Figure 5.5.1. Contour plot of subsurface temperature anomaly (°C) for a salt dome of height 3 km and radius 600 m. Horizontal exaggeration is 2:1 (Modified after O‘Brien and Lerche 1987)

and in the third part particular focus will be placed on the effects of salt-induced temperature disturbances on groundwater flow processes. Numerical examples of thermally-driven brine flow based on the North East German Basin (NEGB) are subsequently described.

5.5.2 Impact of salt structures on temperature field and oil maturation 5.5.2.1 General Concept Salt has a thermal conductivity two to four times greater than that of other sedimentary rocks. Values can be as high as 5 to 6 W m-1 K-1 (Cermak and Rybach 1982; Lerche and O‘Brien 1987). Therefore, a salt dome buried in strata of much lower thermal conductivity will act as a conduit for heat transport vertically and horizontally. This preferential path for heat conduction causes high-temperature anomalies in the surrounding sediments, as reported in many basins (Bayer et al. 1997; Vosteen et al. 2004). The temperature anomaly is defined as the difference between the temperature observed at a point and the regional trend at that subsurface depth. The magnitude of the temperature disturbances depends on the size, shape and depth of the salt diapir as described in detail by O‘Brien and Lerche (1987, 1988) and Yu et al. (1992). Positive and negative anomalies are found respectively above and beneath salt structures. Figure 5.5.1 shows the contours of the temperature anomaly generated by a single salt dome (O‘Brien and Lerche 1987). The largest positive and negative temperature anomalies are to be expected at the top and bottom of the salt dome, while the anomaly goes to zero along a surface which passes through the middle of the dome. The temperature anomaly also extends beyond the edges of the salt dome. Furthermore, there are considerable temperature differences between internal positions in the salt and external positions in the sediments. Sedimentary basins, however, host multiple salt bodies so that their individual thermal effects are mutually interfering and appear as combined temperature pattern (Yu et al. 1992). For example, in the Gulf of Mexico, the anomaly

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Figure 5.5.2. Location of salt diapir Büsum in the Central European Basin System with adjacent Mittelplate exploration wells (lower part), and of salt diapirs and salt pillows in northern Germany (upper part, after Lokhorst et al. 1998)

is as much as 30 °C above salt-bodies and the temperature contrast between internal and external positions is as much as 50 °C (Yu et al. 1992). So far several studies have been published dealing with temperature anomalies around salt structures. Most of them are theoretical approaches, like Jensen (1990); Petersen and Lerche (1996). Empirical studies in basins containing salt lithologies, for example in the area of the Gulf of Mexico, were published by McBride et al. (1998) and O‘Brien and Lerche (1988). Jensen (1983) investigated the temperature field around a salt diapir in Denmark. The published studies in the CEBS are limited to the investigation of the present-day temperature field (Färber 1984) or 2D temperature simulations (Neunzert 1998). Salt temperature effects are particularly relevant for hydrocarbon thermal maturation (Lerche and Lowrie 1992) since oil accumulations are commonly found in association with salt domes, and oil and gas generation react most sensitively to changes of temperature. Furthermore, thermal disturbances related to salt domes will also have an impact on groundwater transport processes. With regard to maturation of organic matter, the strong thermal effects of salt structures influence chemical reaction rates. Thermal maturation in sedimentary environments is a kinetically controlled process that can be

modelled as a first-order chemical reaction. Therefore, as explained by O‘Brien and Lerche (1987) the reaction rates will increase in relation to increased temperature gradients, i.e., on the upper flanks of salt diapirs. On the other hand, on the lower flanks of the salt diapirs, the reaction rates will be lowered. As a result, an enhancement of thermal maturation of any organic matter rich source rock can be expected near the top of the salt. Likewise, negative temperature anomalies will lower or even prevent the maturation of organic matter at the base of the salt diapirs. In general the impact of salt-related thermal disturbances is to enlarge the hydrocarbon window. Clearly, the timing of salt movement and changes in the shape of salt domes over time are very important in order to simulate the correct thermal and maturity history in the vicinity of a salt plug. For this purpose accurate seismic interpretation and structural balancing are prerequisite. The next section illustrates an example based on the Büsum salt diapir (CEBS).

5.5.2.2 The example of the Büsum salt diapir One of the prominent salt structures of the CEBS is the salt diapir Büsum (Fig. 5.5.2). There, a 3D numerical modelling study was performed to investigate the influence of

Temperature fields, petroleum maturation and fluid flow in the vicinity of salt domes

the salt both on the temperature and maturity field of the surrounding sedimentary rocks as well as the temperature and maturity history since the Permian. Calibration of the model was done with measured bottom hole temperatures and vitrinite reflectance data from wells located in the study area. There is a good data base available for this salt diapir because it seals the Mittelplate field which is the largest oil field in Germany. Study area The study area is located in the central part of the North German Basin, in the coastal area of the German North Sea (see Fig. 5.5.2). It covers the salt structure Büsum and its surrounding rim synclines. The salt structure Büsum is a NNE-SSW elongated salt diapir with a northern culmination and salt overhangs surrounding it. It has a length of about 16 km and is 7 km wide. The top of the salt diapir rises from about 1200 m below sea level in the south to about 500 m below sea level in the north. It is one of the double salt walls, which are prominent in the central part of the North German Basin, and is made up of Zechstein and Rotliegend salt with a crestal collapse graben structure at the top of the salt dome. In addition to the salt diapir, there is a salt pillow located to the southeast of the study area, within the Keuper strata. It is assumed that this salt pillow was formed by relocated Permian salt. Located at the western flank of the salt diapir is the biggest German oil field, called Mittelplate, and thus several

exploration and production wells have been drilled in this area, the deepest to a depth of about 3000 m, reaching Liassic strata. Figure 5.5.3 shows a cross section through the salt dome and the adjacent rim synclines. The Jurassic petroleum system comprises the Upper Liassic oil prone source rock (Posidonia Shale), the Middle Jurassic reservoir sandstones and the Middle Jurassic sealing shales. The oil is stratigraphically trapped at the western flank of the salt diapir Büsum (Grassmann et al. 2005). Salt model A 3D model including the salt structure Büsum and its surrounding rim synclines was built (see Fig. 5.5.3). The geometry of the model is based on depth structure maps that were made available by the consortium of RWE Dea AG and Wintershall AG, Germany and completed with depth maps from Baldschuhn et al. (2001). Results from an unpublished gravity modelling study gave the presentday shape of the salt diapir Büsum. The evolution of the salt dome over time was adopted from a 2D structural restoration study while lithology data were adopted from well reports. The 3D model covers an area of about 780 km2. As initial thickness of the Permian salt layer a total thickness of 2500 m was assumed, based on Kockel (1995) and Lokhorst et al. (1998). Based on a mass balance approach palaeo thickness maps of the salt layer were constructed for subsequent time steps, starting from the initial thickFigure 5.5.3. View into the salt diapir Büsum and adjacent rim synclines developped based on seismic interpretation and 3D-PetroMod final element gridding. Yellow: Quaternary and Tertiary, green: Cretaceous, dark blue: Jurassic, violet: Triassic, light blue: Permian

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ness to present-day situation. Best results when modelling complex shapes of salt structures are obtained by using a salt piercing tool, that substitutes selected cells of the 3D mesh at a certain time by a defined lithology including the change of petrophysical properties. Structural balancing for the salt plug Büsum was performed by RWE Dea AG along two seismic lines crossing the salt plug with the software package 2DMove. Furthermore, several depth maps from different horizons are available. The salt plug Büsum is made up of Rotliegend and Zechstein salt. Wells that were drilled in the rim syncline of the salt plug reached their maximum depth in Liassic strata at about 3000 m. Vitrinite reflectance The most important parameter to calibrate simulated thermal histories is the measurement of the thermal maturity of organic matter by means of vitrinite reflectance (VRr) in sediments (Senglaub et al. 2006; Littke et al. this volume). For this purpose dark shales and siltstones were selected for the measurements. The samples were mostly taken from the stratigraphic units Tertiary, Early Cretaceous, Dogger and Liassic up to a depth of about 3000 m (as located in Fig. 5.5.2).

Vitrinites are small particles derived from higher land plants which occur dispersed in almost all sedimentary rocks and in concentrated form in coals. Reflectance of vitrinite has long been known to increase systematically and irreversibly with increasing thermal stress and has therefore been used as coalification parameter for brown coals and bituminous coals (e.g., Taylor et al. 1998). In addition, vitrinite reflectance data have been used more and more to understand the thermal stress and thus the burial history of rocks other than coal. It has, however, to be taken into account that the identification of autochthonous vitrinite is by far more difficult in rocks other than coal (e.g., Radke et al. 1997; Taylor et al. 1998) and that the variability of reflectance data is greater in such rocks (Bostik and Foster 1975; Scheidt and Littke 1989). Nevertheless, vitrinite reflectance has become the primary calibration tool for burial and temperature history modelling since the early work of Lopatin in 1971 published in English by Waples (1980). These authors presented an empirical approach to calculate TTI values (time-temperature index) from temperature histories which could be translated into vitrinite reflectance. These calculated vitrinite reflectance values were then compared to the real data. In case of significant discrepancies, the temperature model had to be modified until a satisfactory fit was achieved. The same principle is still applied in modern basin modelFigure 5.5.4. Vitrinite reflectance data for wells Mittelplate 1 and 2 and calculated trend lines using the algorithm of Sweeney and Burnham (1989)

Temperature fields, petroleum maturation and fluid flow in the vicinity of salt domes

ling studies, but since about 1990, the EASY%Ro algorithm of Sweeney and Burnham (1990) has mainly been used instead of the TTI approach. Vitrinite reflectance measurements were performed on the wells Mittelplate 1, Mittelplate A9, Mittelplate A13, Dieksand 1 and Dieksand 4. For this purpose, a ZEISSPhoto-microscope III with a 40/0.85 oil immersion objective under normal light at a wave length of 546 nm was used. The aim was to make a minimum of 50 single measurements on each sample, as recommended by Barker and Pawlewicz (1993), to obtain sufficient measurement accuracy. A summary of vitrinite reflectance data on samples from the greater study area is found in Rodon and Littke (2005). For the analysed samples, the mean reflectance (VRr) was determined. Most of the samples which provided useful data are from Jurassic and Early Cretaceous strata, from depths between 1980 m and 2680 m. The Dogger samples reach values of 0.44% VRr up to 0.71% VRr. Modelling results Calibration of the models was performed based on an extensive data set on vitrinite reflectance. In figure 5.5.4, this data set is plotted together with calculated vitrinite reflectance-depth trend lines based on the method of Sweeney and Burnham (1990). The 3D-model allowed calculation of the complex temperature and maturity field in the vicinity of the Büsum structure. The salt chimney effect due to the high thermal conductivity of the salt leads to relatively low tem-

peratures below the salt and high temperatures above. The temperature difference compared to the rim synclines can reach up to 17 °C. With respect to maturity, there is also a difference of up to 0.2% VRr observed, with a “too low” maturity below and a “too high” maturity above the salt diapirs. This difference can have a significant effect on petroleum generation and accumulation. For example, the coal-bearing Carboniferous rocks underlying the Permian salt tend to be highly mature in parts of the CEBS. In this pre-Permian source rock sequence, high maturity was already reached in pre-Tertiary times; therefore no significant late gas generation during the Tertiary and Quaternary was possible (Neunzert 1998) leading to gas accumulation. Only further south in the area of Lower Saxony was pre-Tertiary maturation less pronounced, allowing for significant methane generation from Carboniferous source rocks (Littke et al. 1995). There, large gas fields exist, including the giant Groningen field. In the north, however, where Carboniferous source rocks were buried to much greater depth, this high maturity will only be reduced below thick salt domes due to the chimney effect (Fig. 5.5.5). Accordingly, there may be “sweet spots” for petroleum exploration in such areas, where less mature Carboniferous source rocks still had methane generation potential during the Tertiary. A different situation exists with respect to oil generation from the major source rock, which is the Liassic Posidonia Shale (Grassmann et al. 2005). This rock is generally immature to early mature or – even worse – eroded in this area. Increased thermal maturity in the vicinity of the upper parts of the salt diapir could be a positive factor here, although burial depth has the greatest influence on petroleum generation (Rodon and Littke 2005).

Figure 5.5.5. Cross section through the Büsum diapir and adjacent rim synclines with temperature isolines (left) and vitrinite reflectance isolines (right)

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5.5.3 Fluid flow in salt An essential element of a petroleum system in sedimentary basins is the presence of a seal. Rock salt is known to be the best seal for hydrocarbon accumulations, based on three key properties: First, in-situ permeability and porosity of rock salt are very low, even at a burial depth of merely 70 m (Casas and Lowenstein 1989; Urai, Schléder et al. this volume). Second, the near isotropic stress state provides resistance to hydrofracturing. Thus, especially under conditions of extension tectonics, rock salt will hydrofracture under a higher fluid pressure than shale does (Hildenbrand and Urai 2003), because the minimum principal stress (s3) is higher. Third, plastic deformation of rock salt in nature is ductile and therefore non-dilatant (e.g., Ingram and Urai 1999; Popp et al. 2001). This is reflected by Downey‘s (1984) widely accepted ranking of seals: salt → anhydrite → kerogen-rich shale → clay shale → silty shales → carbonate mudstone → chert. On the other hand, under suitable conditions all rocks can lose their sealing capacity. However, the geological conditions for loss of seal capacity of rock salt are not well known.

As a first step, we should differentiate between fluid flow in rock salt (halite) sensu strictu, and in evaporate sequences. Many salt basins are characterised by the occurrence of carbonate rocks within such evaporitic sequences. In the CEBS these carbonates can act as reservoir rocks for oil and gas, proving that active fluid flow has taken place. In eastern Germany, oil reservoirs are present in Zechstein carbonates. Lateral migration over distances of tens of kilometres has probably charged these reservoirs (Fig. 5.5.6; Hindenberg 1999). This self- charging system exists due to the fact that the carbonates formed the marginal facies of the Zechstein basin. At the same time, in more basinal settings, marls and fine grained carbonates were deposited with higher organic carbon content, possessing a source rock potential. Thus intraformational flow within the carbonates towards the most porous and permeable facies filled the reservoirs which were sealed by rock salt. Further towards the west, the same carbonates act as gas reservoirs, but also as gas source rocks. Whereas the bulk of the methane gas in this area is derived from Late Carboniferous coal-bearing rocks (Littke et al. 1995),

Figure 5.5.6. Schematic illustration of facies distribution in Zechstein carbonates of the CEBS (after Hindenberg 1999) including source and reservoir rock facies

Temperature fields, petroleum maturation and fluid flow in the vicinity of salt domes

a minor contribution has to be attributed to a carbonate source rock. As these gas pools with “carbonate” signature mainly occur in Rotliegend reservoirs juxtaposed to the Zechstein sequence, it is probable that migration of gas out of the carbonates into the reservoirs has occurred (Littke et al. 1996). In the Lower Saxony Basin constituting the southern part of the CEBS, carbonate stringers are partly in contact with the underlying older strata and not underlain by salt. Some of these carbonates act as gas reservoirs, having vastly different contents of methane and carbon dioxide. In this case, the structural setting seems to be decisive for the quality of the gas filling. In particular, early filling by methane gas from underlying Carboniferous source rocks followed by decoupling from later inflow of carbon dioxide seems to be a prerequisite for the existence of high quality gases (Petmecky 1998; Petmecky et al. 1999; Krooss et al. this volume).

Observations above clearly demonstrate that there is fluid flow taking place into evaporites, out of evaporites and inside evaporites. Furthermore, simple mass balances on total gas generation in the thick Late Carboniferous source rock sequence of the southwestern CEBS prove that much more gas was generated post-Zechstein than is presently trapped (Littke et al. 1995). Accordingly, this gas must have migrated out of the Palaeozoic sequence through the Zechstein evaporites towards the surface. Nollet et al. (2005) studied vein systems in the Buntsandstein of the CEBS. Based on microstructural and geochemical arguments they proposed a regional high pressure cell during burial of this sequence, with the origin of the fluids in the Zechstein, which must have migrated through the rock salt sequence into the Bunter.

Figure 5.5.7. Macroscopic (left) and microscopic (right) appearance of solid bitumen in rock salt (Ara salt, Oman; see Schoenherr et al. 2007a,b for details)

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These results imply that large scale fluid migration through the evaporitic sequence can be quite common. In-situ permeability of undisturbed rock salt is ∼ 10-21 m2 (Bredehoeft 1988; Peach and Spiers 1996; Popp et al. 2001). This low permeability allows rock salt to seal large hydrocarbon columns and fluid pressure cells over geological time. Two processes are known to increase permeability. The first is microcracking and associated dilation (Peach and Spiers 1996; Popp et al. 2001), and the second is the formation of topologically connected brine-filled pores and triple-junction tubes in halite grain aggregates at a pressure and temperature corresponding to depths > 3 km (Lewis and Holness 1996). In the triaxial deformation experiments of Lux (2005), rock salt became permeable by the formation of grain boundary cracks at low rates of fluid-pressure increase. If the fluid-pressure (P f) is increased rapidly to a sufficiently high excess pressure (Pf > σ3), rock salt leaks by hydrofracturing. Peach and Spiers (1996) proposed that this process could also occur under low effective stress during natural deformation of rock salt at great depth and high pore-fluid pressures. More theoretical background on permeability and microfractures in rock salt is discussed in Schoenherr et al. (2007a) and Urai, Schléder et al. (this volume). One of the best known examples of oil flow in rock salt is found in the Infra-Cambrian Ara Group of South Oman (Schoenherr et al. 2007b). The Ara Group consists of marine platform sediments, representing at least six third-order cycles of carbonate-evaporite sedimentation, of up to 4 km thickness (Mattes and Conway Morris 1990). Each cycle is characterised by sedimentation of up to 1000 m Ara Salt at very shallow water depths, followed by the deposition of 20 to 250 m thick isolated carbonate platforms (the so-called “stringers”) during transgressive periods. Most of the carbonate stringers contain fluids at very high overpressures. The carbonate stringers have undergone intense diagenetic modifications with locally extensive cementation by halite and solid bitumen (Schoenherr et al. 2007a). They have been interpreted to be a self-charging system with first oil charge from mature Type I/II source rocks (Peters et al. 2003) during the early Cambrian to Ordovician at maximum burial temperatures after deposition of the overlying Haima Supergroup (Visser 1991; Terken et al. 2001). In addition, geochemical studies suggest external (pre-Ara) oil and gas charge (Al-Siyabi 2005). Presence of solid bitumen as a relic of former oil is evident from macroscopic observations not only in the carbonates (where it may be expected), but also in the rock salt (Fig. 5.5.7, left; Schoenherr et al. 2007a). Microscopy revealed that solid bitumen is both present on grain boundaries and in microcracks (Fig. 5.5.7, right). The black rock

salt cores clearly indicate that distinct parts of the Ara Salt lost sealing capacity for oil. Oil initially leaked into the salt when oil pressure in the reservoir exceeded the nearlithostatic fluid pressures of the Ara Salt by a few tenths of a MPa to overcome the capillary entry pressure, followed by diffuse dilation and marked increase in permeability, in agreement with laboratory-calibrated dilation criteria of rock salt. Oil could flow into the salt as long as the oil pressure remained larger than the minimum principal stress in the rock salt. A detailed explanation is given in Schoenherrr et al. (2007a,b). The conditions inferred from this study represent the ultimate sealing potential of rock salt in the deep subsurface, and can be applied to rock salt seals in any tectonic setting. Another study of fluid flow in salt is presented in by Schléder and Urai (2005), see also Urai, Schléder et al. (this volume). In summary, lithostatic fluid pressures increase the permeability of rock salt by many orders of magnitude, allowing flow of water and hydrocarbons until the fluid pressures decrease. There is increasing evidence that this has occurred in many salt-bearing basins, providing an explanation for many observations, and an indication for palaeo-overpressures.

5.5.4 Impact of salt structures on groundwater transport processes within sedimentary basins In this paragraph, the effects of salt-induced temperature disturbances on groundwater transport are described. For this purpose, a brief overview of the main driving forces of groundwater flow in sedimentary basins is given. Then numerical examples based on the North East German Basin (NEGB) will serve to illustrate the principle impact of salt domes on thermally-driven brine flow.

5.5.4.1 Brief description of driving forces in large-scale groundwater flow systems Sedimentary basins are subjected to several forces known to cause large-scale groundwater migration, each characterised by a typical flow rate as reviewed by Bjørlykke et al. (1988); Garven (1995); Person et al. (1996) and Ingebritsen and Sanford (1998). Topography-driven flow (Fig. 5.5.8A) is the dominant regional-scale groundwater flow in sedimentary basins both in the shallow and deep sub-surface (Freeze and

Temperature fields, petroleum maturation and fluid flow in the vicinity of salt domes Figure 5.5.8. A Topography driven flow in an uplifted sedimentary basin, B Thermally induced flow in a geothermal basin and C Geologic forcing, flow due to tectonic compression (Modified after Garven 1995)

Witherspoon 1967). This regional flow occurs when differences in the hydrostatic head drive fluid from high-elevation recharge areas to low-elevation discharge areas. Flow lines and rates depend on several factors as the geometry of the aquifers and their physical properties (e.g., hydraulic permeability). Typical maximum flow rates range from 1 to10 m year-1. In sedimentary basins, the presence of chemical compounds and/or the geothermal field induce fluid-density variations which in turn drive groundwater flow. Fluid motion caused by density difference due to temperature variations is called free convection (Fig. 5.5.8B) while gravitational convection is the term used when the convective currents are induced by variations in solute concentration.

Ongoing geological processes such as sediment compaction, tectonic compression (Fig. 5.5.8C), hydrocarbon generation or degassing of magma can induce significant fluid flow. These processes are referred to as geological forcing (Neuzil 1995) and their flow rates span a large range of velocities. For example, compactionally driven pore-water flow rates are usually very slow, in the order of 10-6 to 10-3 m year-1, while tectonic compression can induce flow rates of 0.5 m year-1 (Garven 1995). These processes are coupled! An important feature of deep groundwater flow within sedimentary basins is that mechanical, hydrological, thermal, and chemical mass transfer processes are fully coupled.

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Figure 5.5.9. Top thermally induced brine plumes developing from a deep salt sheet. Bottom brine lenses induced by an instable density stratification and thermohaline convection. The bold vectors indicate the direction of the cellular motion, the dashed lines are the isotherms in °C

Transport processes take place over large temporal scales during which basin deformations are often substantial. Therefore, there is a mechanical coupling between the driving forces and the structural evolution of the basin. Furthermore, increases in sub-surface fluid pressures induce rock dilation and porosity increase. In contrast, mineral precipitation reduces rock porosity, decreasing groundwater flow rates (Garven 1995). Over the large spatial scales encountered in geothermal basins, temperature and solute concentration strongly vary. These variations modify the physical properties of both fluid (density, viscosity) and rock units (hydraulic conductivity, porosity) leading to new dynamic effects of the system. The archetypal example for sedimentary basins is the coupling of heat and dissolved halite which have different rates of diffusion (heat diffuses faster than salt): the flow

is then called thermohaline convection (Nield and Bejan 1999). Since in geothermal basins the temperature gradients increase with depth, heat acts as a destabilising potential. Two major scenarios for thermohaline convection to occur can be distinguished. (i) When salt concentration increases with depth (salinity is stabilising), the deeper brines are heated from below and therefore less dense: an upward flow is triggered leading to the formation of convective cells (Fig. 5.5.9 top). As the flow progresses, the brines will cool off quickly while loosing little salt diminishing the upward buoyancy force. Eventually, brine will start sinking. (ii) Another possible scenario is when salinity gradients also act as destabilising factors. That is the case when brine forms in shallow areas of the basin (e.g., from shallow salt structures). The denser fluid will therefore sink into the deeper and hotter part of the basin. At the same time, lighter and hotter fluids will move upward

Temperature fields, petroleum maturation and fluid flow in the vicinity of salt domes

Box 5.5.1 Mathematical formulation of the thermohaline flow problem The governing equations of thermohaline convection in a saturated porous media are derived from the conservation principles of linear momentum, mass and energy (e.g., Kolditz et al. 1998; Nield and Bejan 1999). They are briefly reported here by the following set of differential equations:

(5.5.1)



(5.5.2)



(5.5.3)



(5.5.4)

Equation 5.5.1 is the equation of fluid mass conservation. S0 is the medium storativity which physically represents the volume of water released from (or added to) storage in the aquifer per unit volume of aquifer and per unit decline (or rise) of head φ . QBoussinesq is the Boussinesq term which incorporates first order derivatives of mass-dependent and temperature-dependent compression effects. q is the Darcy (or volumetric flux density velocity) defining the specific discharge of the fluid. Darcy’s law is expressed by equation 5.5.2 where K is the hydraulic conductivity tensor, equation 5.5.3 is the equation of solute mass conservation where φ is the porosity of the porous medium, C is the mass concentration, D is the tensor of hydrodynamic dispersion and QC is a mass supply. Equation 5.5.4 is the energy balance equation of the fluid and porous media. cf and cs are the heat capacity of the fluid and solid respectively, T is the temperature, λ is the tensor of hydrodynamic thermodispersion. Constitutive and phenomenological relations of the different physical parameters involved in the equations are needed to close this coupled system. Here the hydraulic conductivity relation and the equation of state (EOS) for the fluid density are given:

(5.5.5)



(5.5.6)

The hydraulic conductivity tensor K is related to the reference fluid density . g is the gravitational acceleration, k is the tensor of permeability, μf (C,T) takes into account the fluid viscosity effects due to temperature and concentration variations. The EOS for the fluid density (Eq. 5.5.6) is related to the reference temperature T0, pressure p0 and concentration C0. is the mass concentration ratio,

being the fluid density at saturation. β (T,p) is the coefficient of thermal

expansion and γ (T,p) is the coefficient of compressibility. The flow and transport equations (Eq. 5.5.2, 5.5.3, 5.5.4) for thermohaline convection are non-linear and strongly coupled since temperature and salinity control the fluid density ρf and dynamic viscosity μf (C,T). The variation of fluid density is essential for the modelling of thermohaline convection because of its primary importance for calculating the correct buoyant force included in the equation of motion (i.e., generalised Darcy’s law Eq. 5.5.2). Fitted polynomial expressions are commonly used for temperature, pressure and salinity dependences of the fluid density (Sorey 1976).

.

Stability criteria

A dimensional analysis of the governing balance equations (Eq. 5.5.1 to 5.5.4) allows the definition of several adimensional numbers (Nield and Bejan 1999). The key dimensionless number is the Rayleigh number (Ra), which is the ratio between buoyancy-driven forces and resisting forces caused by diffusion and dispersion: Thermal Rayleigh number Rat

(5.5.7)

Solutal Rayleigh number Ras

(5.5.8)

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where K is the hydraulic conductivity as defined in equation 5.5.5, α introduces the effect of a density change due to the concentration of the solute at temperature and pressure. β is the coefficient of thermal expansion at constant pressure and concentration, Λ is the thermal diffusivity, Δ C and Δ T are respectively the concentration and temperature variation, d is a characteristic length of the porous media (e.g., the layer thickness), ε is the porosity, Dd is the coefficient of molecular diffusion. The solutal and thermal Rayleigh numbers are related by:

(5.5.9)

where the dimensionless numbers in connection with heat and mass transport are: Buoyancy ratio (Turner) N is the relative effects of concentration and temperature on controlling groundwater density



(5.5.10)

Lewis number Le is the ratio of thermal to mass diffusivity

(5.5.11)

The stability criteria are thoroughly explained in Nield (1999), Diersch and Kolditz (2002). Here the main points are recalled.

The monotonic instability (or stationary convection) boundary is a straight line defined by Ra = Ra + Ra , where Ra is the · critical Rayleigh number. The critical Rayleigh number defines the transition between dispersive/diffusive solute transport C

T

S

C

(Ra < RaC) and convective transport by density-driven fingers (Ra > RaC). RaC depends on the boundary conditions, geometry and anisotropy (Nield and Bejan 1968).

· The region delimited by is a stable regime characterised by pure conduction and no convection. between steady state convective cells develop as two-dimensional rolls rotating in a clockwise or counter-clock· Inwisea range direction. A second critical Rayleigh number Ra = 240 – 300 is identified as an upper limit. For Ra + Ra > Ra the convection regime is unstable and characterised by a transition to an oscillatory and transient con· vection behaviour. c2

T

s

c2

From equations 5.5.7 and 5.5.8, it can be seen that given the diffusivities of the unit, Rayleigh numbers are directly proportional to the thickness of the unit and hydraulic conductivity. Therefore convective flows (i.e., high Rayleigh numbers) probably occur within thicker and more permeable layers. Figure 1. Schematic temperature and concentration profiles of a homogeneous porous layer heated from below. Given the physical properties of the medium, the Rayleigh numbers can be determined (Eq.5.5.7 and Eq.5.5.8)

owing to the thermally induced buoyant forces. Consequently, brine lenses and convection cells form (Fig. 5.5.9 bottom). Thermohaline convection can develop cells at flow rates approaching 1 m yr-1 (Evans and Nunn 1989; Garven 1995) which are strong enough to control temperature and concentration fields. On the other hand, vigorous topographically-driven groundwater flow can overwhelm free convection and modify the thermal structure of the basin (forced convection). Precisely, it causes cooling in recharge areas and increase heat flow in discharge areas. Temperature differences up to 50 °C can be expected between recharge and discharge areas

(Ingebritsen and Sanford 1998). In some cases, thermally-induced convection and topography-driven flow coexist (Raffensperger and Garven 1995a; Thornton and Wilson 2007) leading to a mixed convection regime. The onset of thermohaline convection can be derived numerically. The mathematical formulation and the stability criteria for the thermohaline problem are given in Box 1. A key number which controls the flow dynamics is the Rayleigh number. If the Rayleigh number is large enough, then cellular motion can develop. As a result, a multitude of stability analyses based on laboratory experiments were carried out on saturated porous media with vertical gradients of temperature and salinity in order

Temperature fields, petroleum maturation and fluid flow in the vicinity of salt domes

to determine the critical Rayleigh number for the onset of thermohaline convection (Horton and Rogers 1945; Lapwood 1948; Elder 1967; Trevisan and Bejan 1987; Kubitschek and Weidman 2003). Comment on the drawbacks of adimensional analysis for real basin systems: While Rayleigh theory can be successfully used for the study of convective flow in homogenous systems, its applicability for transport processes within real sedimentary systems is seriously questioned (Raffensperger and Garven 1995a; Simmons et al. 2001). This criterion assumes that a steady-state flow takes place in a homogeneous system under steady state conditions. A Rayleigh number may be derived for heterogeneous systems provided that the properties vary within only one order of magnitude (Nield 1994). In real-site applications, this is often not the case: the physical parameters of the basin are subject to large heterogeneities. Furthermore, stability analysis based on dimensionless numbers involves the definition of a characteristic length-scale of the porous media (d Box 5.5.1 Fig. 1) through which temperature and solute gradients are supposed to vary linearly. These conditions are not satisfied in any transport processes occurring within real basin systems. Usually, concentration and temperature gradients are non-linear due to salt structure disturbances. The thicknesses of the stratigraphic units of sedimentary basins can vary from a few meters to several kilometers because of the strong salt tectonics. Therefore, the definition of a representative length scale is rather problematic. From these considerations, it follows that the Rayleigh criteria are not appropriate (and perhaps even impossible) for analysing coupled transport processes within real system.

Clearly, any effect which causes a significant variation in temperature gradient will play a fundamental role in determining the groundwater regime. Therefore, owing to their physical properties, salt diapirs provide a unique geological environment for thermohaline convection to occur. In recent decades, the perturbation of the geothermal field associated with salt diapirism has been studied for thermally-induced fluid convection (e.g., Fogg et al. 1983; Hanor 1987; Ranganathan and Hanor 1988; Evans and Nunn 1989; Evans et al. 1991; Thornton and Wilson 2007). There are a multitude of environmental circumstances where thermohaline convection can arise, includ-

Salt diapirs: a unique geological environment for thermohaline convection The relative importance of the above mentioned coupled driving forces on fluid flow varies depending on the tectonic and lithologic conditions. It is proven that thermallyinduced flow can lead to transport of dissolved compounds over large spatial scales and in significantly shorter times compared with diffusion alone (Hanor 1987; Diersch and Kolditz 2002). Furthermore, it has been shown that thermally-induced flow is a major factor controlling deposition of minerals in basins and, therefore, affecting basins diagenesis. For instance, numerical investigations of coupled transport processes showed that free convection is responsible for ore deposits in the Mc Arthur Basin, Australia (e.g., Garven et al. 2001; Yang et al. 2004a,b) and in the Athabasca Basin, Canada (Raffensperger and Garven 1995a,b).

Figure 5.5.10. Top mass streams lines. Bottom isotherms and concentration contour (Modified after Evans and Nunn 1989)

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ing transport of pollutants released from waste disposal in salt rock formation (Evans and Nunn 1989), or salt layers embedded in aquifers (Sarkar et al. 1995) can be mentioned. The Gulf Coast region of the United States is one of many examples and was one of the first case studies of thermohaline convection in salt-dome environments.

5.5.4.2 Example: Gulf Coast region of the United States The salt domes of the Gulf Coast have been thoroughly explored in the search for oil and gas (Mace et al. 2006). Hydrochemical investigations showed that fluids around several salt domes have anomalous salinities owing to the local dissolution of halite from the salt flanks

(Workman and Hanor 1985; Bennet and Hanor 1987; Enos and Kyle 2002). Groundwater that is in contact with the salt dome reaches concentrations up to saturation level (i.e., 345 g/L). Plumes of dense saline water are gravitationally unstable and have the potential to sink. Furthermore, hydrochemical patterns indicate that warm fluids from deep formations discharge into shallower units. The coupling of these downward and upward migrating fluids is likely to form convective flow systems of 1 km diameter. A comprehensive review of fluid-rock interaction in salt-dome environments for this area is given by Posey and Kyle (1988). Numerical models were built in order to investigate the feasibility of these flows. According to the Ranganathan and Hanor models (1988) salinity-driven convection cells

Figure 5.5.11. The NEGB study area: location, salty water distribution (after Grube et al. 2000) and sea grass plants

Temperature fields, petroleum maturation and fluid flow in the vicinity of salt domes

achieve maximum flow velocities of 0.01 to 1 m yr-1. Further modelling showed that the temperature disturbances caused by the presence of a salt structure can set up buoyancy forces which have the potential to drive large-scale brine convection (Evans and Nunn 1989; Evans et al. 1991). In their calculation, two distinct kilometre-scale convective cells develop: one cell near the top of the dome with upward flow at the salt edge, and another extending to the bottom of the system with downward flow (Fig. 5.5.10 top). A salt plume extends away from the salt dome as a result of upward flow (Fig. 5.5.10 bottom). In summary, the disturbed thermal regime around the dome generates an instable fluid-density stratification which induces upward brine movement along salt flanks. Eventually, warm fluids discharge into the shallower units. However, these models were based on a single salt dome surrounded by homogenous sediments while in sedimentary basins different salt bodies are in close juxtaposition and cut through heterogeneous units. More detailed models with regard to the basin lithology and internal structures are needed in order to understand the geological conditions in which thermally-induced convective flow occurs.

5.5.4.3 Numerical example of thermallyinduced flow in relation to salt dome environment (including chemical observations): the N-E German Basin An interesting example illustrating thermally-driven processes in a sedimentary basin hosting many salt structures

is provided by the North-East German Basin (NEGB) as part of the Central European Basin System (CEBS) as described in detail by Magri (2005); Magri et al. (2005a,b, 2007). Salty-water evidences in the NEGB Extensive evidence for rising saline waters has been gathered in different regions of this basin (Fig. 5.5.11). It can be seen that the main brine plume of saline groundwater slants across the basin stretching over 250 km from the western part to the south-eastern area near Berlin. In addition, isolated brine formations with an extension in the order of 20 km can be observed in the northern and western part of the domain. The discharge area of brackish waters is occurring in the south-eastern part of the NEGB. Precisely, the majority of the saline springs are observed in the neighbourhood of the Elbe and Havel rivers and to the south of Berlin. Only a few are encountered away from these locations. Additional evidence of salty groundwater is given by plants commonly found along sea beaches or in salty soils, such as seashore salt grass, which grow in different areas of the basin (Fig. 5.5.11). The photo was taken in 2004 during a field trip to Gröben, 50 km south of Berlin. The spontaneous growth of seashore grass far from the Baltic Sea coast is unusual and signals the existence of highly salty soils in the inner part of the basin. Although this phenomenon has been observed for decades (Heck 1931; Johannsen 1954; Schirrmeister 1996; Hannemann and Schirrmeister 1998; Grube et al. 1996, 2000), an unsolved feature of these saline springs is their Figure 5.5.12. Left: Zechstein depth map; Right: NEGB surface elevation together with location of the observed springs (Schirrmeister 1996) and salty groundwater interface

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temporal and spatial instability. Furthermore, the observed brines at the surface are a few degrees hotter than the average fluid temperature in the soil. Figure 5.5.12 (left), illustrates the occurrence of shallow salt domes (dark patches) and thin deeper salt pillows (light grey areas) together with the salty groundwater distribution. It can be seen that no obvious spatial correlation between salt structures and near surface salt occurrences exists. Brine patterns can also be found far away from salt dome crests. Most of the salty groundwater stretches are in the lowland along the rivers system (Fig. 5.5.12, right). It can be inferred from this observation that the regional flow induced by topography variation plays an important role in driving solutes. Nevertheless, some salty plumes can also be observed in regions far from this area indicating that regional groundwater flow is not the only process transporting solute within the basin. Furthermore, the topography variation is not great enough to lift deep-seated waters up to the surface. Hydrochemical patterns in groundwater provide information about possible salinity sources and flow direction. The hydrochemistry of deep saline waters in the Northern German Basin was investigated amongst many others by Lehmann (1974a,b); Voigt (1977); Hannemann and Schirrmeister (1998); Tesmer et al. (2007); Möller et al. (2007b). The main source of salinity is halite dissolution from the salt domes. Salt content increases with depth until it reaches a saturation point estimated around 345 g/L. The influence of a topography-induced flow regime on salt migration is supported by isotopic analysis which indicate that the regional groundwater flow affects the water cycle down to a depth of at least -500 m (Tesmer et al. 2007). On the other hand, the Rare Earth elements and Yttrium (REY) patterns highlight the existence of upward

directed inter-aquifer flow even at a depth below -1500 m (Tesmer et al. 2007; Möller et al. 2007b). Therefore, intrinsic basin-system mechanisms must exist to overcome the gravity field that would keep dense salt-laden waters in deep seated aquifers. Hydrostatic pressures rise in all stratigraphic units above the Zechstein, suggesting that ascending flows are not due to the existence of any overpressured aquifers. A possible cause could be thermallyinduced flow as indicated by fluid density studies. Fluid density analyses show that from the point of saturation the density decreases with increasing depth. The inversion of the density/depth trend is due to the thermal expansion of the fluid. Precisely, at the saturation depth, temperature effects are dominating and increase the brine volume which leads to a decrease of its density. Within the Cretaceous Keuper interval this effect occurs at depths below 3500 m. Accordingly, the fluid density stratification is unstable and promotes thermally induced convective flows. At this state the question arises, how do saltinduced temperature disturbances drive brine flow? The numerical simulations presented in the next paragraph will highlight the role these temperature disturbances play in controlling upward solute transport The numerical model Large scale simulation of coupled fluid flow, mass and heat transport based on a real geothermal system requires a proper aquifer and fluid model. The aquifer model should include the structural characteristics of the aquifers as well as the physical parameters such as porosities, hydraulic permeabilities, heat conductivity and heat capacity. The fluid model should take into account the chemical and physical fluid characteristics and their spatial distribution, water salinity, chemical components and fluid density.

Figure 5.5.13. Stratigraphic units of the 2D cross-section. The stratigraphic unit abbreviations are: Cz Cenozoic; K1 Late Cretaceous; K2 Lower Cretaceous; J Jurassic; T2-3 Late Triassic; T2 Middle Triassic; Lower Triassic; T1 Buntsandstein

Temperature fields, petroleum maturation and fluid flow in the vicinity of salt domes

Figure 5.5.14. A: temperature profile (°C) and B: Concentration (g/L) resulting from transient free thermohaline convection (t=200 ka)

Aquifer model The incorporated geological data are derived from a three-dimensional structural model of the NEGB (Scheck 1997; Scheck and Bayer 1999). The area covered by the model is approximately 230 x 330 km across and 5 km in depth, consisting of 9 layers of sedimentary fill, including the basement. Figure 5.5.13 illustrates the geological structures of the cross-section used for the numerical simulations. The NEGB is affected by intense salt tectonics. Thick salt diapirs pierce more than 4 km of overlying Mesozoic and Cenozoic strata. Salt crests can also be found 500 m below the surface level. Therefore, depth and thickness of sediment sequences vary greatly within the basin. The physical properties considered within each layer are constant. This first rough aquifer model differentiates only the stratigraphic layers of the model without any spatial variation. More details concerning the physical parameters of the sedimentary layers can be found in Magri et al. (2005a). Local faults are not included. Fluid model The hydrochemical investigations allow us to make some preliminary assumptions for the fluid model. The brine can be considered pure NaCl solution resulting from halite dissolution. For the brines within the Mesozoic strata this assumption provides a sufficient approximation. The saturation concentration of the fluid is reached at 345 g/L of dissolved halite which corresponds to a brine density of 1220 g/L. In geothermal systems, the influence of pres-

sure, temperature and concentration on the fluid density cannot be neglected. Two polynomial expressions which accurately represent the coefficient of thermal expansion β (T,p) and compressibility γ (T,p) for the fluid density (Eq. 5.5.6, Box 5.5.1) have been derived and coded as an extension to the simulation program. A detailed description of these polynomial functions and the implemented code is reported in Magri et al. (2005a). Modelling approach The strongly coupled non-linear equations governing thermohaline convection in porous media (Box 5.5.1) are solved by the use of the commercial Finite Element (FE) program FEFLOW (Finite Element subsurface FLOW system), WASY GmbH. The resolution of the finite element grid is variable and allows modelling of variations in fluid-density. More details, also on boundary conditions, can be found in Magri et al. (2005a). Role of salt-induced temperature disturbances on groundwater transport processes: simulation results Temperature disturbances on brine flow: free thermohaline convection Figure 5.5.14(AB) shows the calculated transient temperature and brine distribution. The principle result is that

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F. Magri · R. Littke · S. Rodon · U. Bayer · J.L. Urai Figure 5.5.15. Zoom of thermohaline simulation results in the salt dome environment. A) Pore water velocity field in mur-1. Pore vector linearly scaled to the largest flow arrow. B) temperature distribution in °C

a disturbed profile develops throughout the sediment fill above the Zechstein unit. The oscillatory pattern is characteristic of a multicellular convective regime. The waves are non-periodic and their amplitudes decrease with depth. In the western part of the basin, the temperature isopleths are nearly flat presenting only one anomaly near the salt dome. In the eastern part of the basin, a long way from the salt diapir, the temperature distribution is less disturbed. Below the Zechstein unit the temperature profile is conductive everywhere, showing the well-known thermal anomalies within the salt diapirs (i.e., concave temperature isopleths). Owing to the strong coupling between heat and mass transfer in thermohaline convection, the saline and thermal plumes developed together spreading over the same areas within the sediment fill (compare Fig. 5.5.14A and B). Above the Muschelkalk, brine plumes develop rap-

idly and penetrate the overburden (Fig. 5.5.14A). Brine fingers form and extend vertically over 3 km throughout the sediments. Salty waters with more than 1 g/L of dissolved halite spread over the surface driven by the disturbed geothermal field. At the western part of the basin, the finger regime disappears. The concentration isopleths are flat and the salt content increases almost linearly with depth. The layered brine stratification is probably a boundary effect. Since the lateral boundaries of the model are closed to fluid flow, mass and heat transfer, the dissolved halite cannot diffuse away in the western direction. Moreover, salt diffusion in the eastern direction is also prevented due to the presence of shallow salt diapirs. Therefore, as time progresses, the sediments are filled with halite. The finger regime evolves into a layered system in which concentration increases with depth. At the eastern ending of the

Temperature fields, petroleum maturation and fluid flow in the vicinity of salt domes

Figure 5.5.16. A: Concentration (g/L) and B temperature profile (°C) resulting from transient mixed convection (t=200 ka). The bold vectors schematically represent the direction of the topography-induced flow

profile, brine plumes do not develop since at that location the temperature gradient is not disturbed. On the other hand, temperature disturbances play a dominant role on brine migration especially in the neighbourhood of salt diapirs. A zoom of the calculated pore velocity and temperature fields in a salt dome environment is shown in figure 5.5.15A. Downward forces resulting from the gravitational field control groundwater flow along salt diapir flanks. The salt-laden water sinks at approximately 1.5 cm yr-1 (Fig. 5.5.15A). By contrast, an upward flow paralleling this descending flow occurs in the overlying unit at approximately 1 cm yr-1. This phenomenon can be explained by the temperature distribution (Fig. 5.5.15B). Because of the thermal conductivity contrast between salt and overlying sediments the isotherms are convex near the edge of the salt diapir. The increased temperature gradient causes a decrease in fluid density near the salt dome. This drives the groundwater flow toward the salt dome and initiates the uprising circulation of brine within the neighbouring sediments. Temperature disturbances and regional flow: Mixed convection When an external factor such as head-driven groundwater flow (forced convection) is imposed on a free thermohaline system the resulting regime is referred to as mixed convection (see brief description of driving forces). Here

the head level is set equal to the topographic relief so that regional and thermally induced flows occur together. In this section, the results from free thermohaline and mixed convection are compared in order to study the interaction between temperature disturbances and the regional flow. Temperature and mass distribution resulting from mixed convection (Fig. 5.5.16A-B, respectively) show significant differences with regard to the profile derived from the free thermohaline simulation (Fig. 5.5.14AB). In the mixed convection regime, the thermally induced brine patterns are affected by the regional flow: the short wavelengths of the temperature oscillations characterising the free thermohaline regime (Fig. 5.5.14B) are not preserved in the mixed convection system (Fig. 5.5.16B). The isotherms are shaped by the regional flow: in the recharge areas, the infiltration of cooler water decreases the temperature gradient whereas uprising of warmer plumes occurs in relation to the discharge areas. As a result, the narrow salty fingers observed in the free thermohaline regime (Fig. 5.5.14A) evolve into a smaller number of larger brine plumes which reach the surface at the discharge areas (Fig. 5.5.16A). The brine patterns display a truncated profile in direct relation to the downward flow of freshwater. Clearly, temperature disturbances and regional flow strongly interact. This interaction can modify the geo-

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Figure 5.5.17. Concentration distribution (g/L) at the end of the simulation resulting from coupled fluid flow and mass transport (i.e., no geothermal gradient)

thermal field and therefore thermally-induced flow. Precisely, decreased concentration and temperature gradients occur in relation to inflow of freshwater while thermally-induced brines are advected toward discharge areas. Importance of salt-induced temperature disturbances in controlling deep groundwater transport processes. Here the geothermal gradient of the upper crust in the NEGB has been excluded from the simulations. This simulation allows us to demonstrate how important salt-induced temperature disturbances are in controlling transport processes.

Figure 5.5.17 illustrates the resulting mass patterns at the end of the simulation run. Within the whole Buntsandstein unit the concentration values range from 200 g/L at its surface to 345 g/L, which is the halite fluid saturation, at the top salt. Above the Muschelkalk, dissolved halite flows in the neighbourhood of salt domes. Layered brine plumes spread at both sides of the salt diapirs and do not stretch vertically. The salt dome environment is hydrochemically unique, i.e., groundwater is subjected to large lateral salinity gradients. Along salt dome flanks the density gradient drives Figure 5.5.18. Coupled fluid flow and heat transport simulation. A Calculated temperature profiles in °C. B Zoom of the thermally induced plumes without vertical exaggeration

Temperature fields, petroleum maturation and fluid flow in the vicinity of salt domes

groundwater flow laterally into the basin. The dissolved halite spreads away from salt bodies by diffusion and rapidly increases the salinity of the pore water throughout the sediments. The resulting brine plumes extend over 30 km in lateral direction. As a result, dissolved halite concentration around salt domes increases with depth. Above salt diapir crests, brackish water with 1 g/L of dissolved salt can reach a depth of between one and half a kilometre. Away from salt domes, the salty plumes do not develop as is the case in the central and eastern part of the basin. The temperature disturbances illustrated so far were generated for deep salt diapirs. However, salt diapir crests are often close to the surface. Steep salt diapirs close to the basin surface provide a source of high salinity for shallow groundwater which can affect the temperature oscillations. In the next section, a numerical simulation will provide new insights into the effects of temperature disturbances on groundwater transport processes in a shallow salt diapir environment. Temperature disturbances and groundwater transport processes in a shallow salt diapir environment The temperature field calculated from a coupled fluid flow and heat transport simulation is illustrated in figure 5.5.18. This structural profile includes a steep salt diapir piercing the sediments up to the surface. Different regimes developed within the profile. Within the salt unit, the thermal regime is conductive. Owing to the strong contrast between the thermal conductivity of the salt and the neighbouring sediments, concave isotherms are found within the salt diapir while convex isotherms are found adjacent to the salt flank.

Above the salt, advective, convective and conductive heat flow affect the whole profile. Within the post-Palaeogene units, the regional flow is dominant. Increased temperature gradients are found in direct association with discharge areas. For instance, the fluid temperature within the Quaternary channel increases by about 2 °C. A similar temperature increase can be observed at the discharge area of the Hemmelsdorf Basin. On the other hand, inflow of cooler water decreases the temperature in the recharge areas. According to the Rayleigh theory in a porous media, the onset of multi-cellular convection is favoured in thick and permeable units (Nield 1968). This is the case for the Palaeogene and the Cretaceous units where a thermally induced convective regime controls the flow. Thermal plumes of 1.5 km height rise vertically from the Cretaceous basis up to the surface, bounded by the regional flow. A zoom of an ascending thermal plume is shown in figure 5.5.18B. The cell radius is 1 km and the flow rate in the central part of the plume is few millimetres per day. In the deeper units the isotherms are not perturbed and the regime is conductive. Temperature and salinity profiles are shown in figure 5.5.19. Highly saline brines protruding from the salt diapir into the Cretaceous overwhelm the less intense thermal convective regime. Heat plumes do not stretch vertically but develop almost horizontally in the brine flow direction (compare Fig. 5.5.19 and 5.5.18). Therefore, the temperature gradient increases horizontally from the salt flank toward the center of the profile. As a result, the temperature field can undergo several inversions with increasing depth in the western part of the profile. In the Eastern part of the basin, thermohaline convection persists within the upper units. Above the horizon-

Figure 5.5.19. Thermohaline simulation. Calculated mass (filled patterns, g/L) and temperature profiles (red dashed lines, °C)

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tally stretched plume, the temperature oscillations generate small convective brine cells (half kilometre radius). As a result, thermally driven saline waters ascend up to the shallow aquifer and spread locally at several points of the surface. On the other hand, in the other units of the profile, the geothermal gradient do not significantly influence brine patterns as explained in details in Magri et al. (in print). In summary, the numerical models have shown that thermally-induced flow is an important process in saltbearing basins, and strongly controls both temperature

and concentration gradients. Owing to the presence of thick salt structures, the geothermal field is disturbed. The salt-induced thermal disturbances in turn induce convection of deep brines. However, it is not the only process. Topography-driven flow also influences the geothermal field and can significantly contribute to brine migration. Whereas the principal effects of thermohaline convection could be shown, much more detailed knowledge of transport properties (hydraulic permeability, thermal conductivity) and their regional distribution, including faults and fractures, are necessary in order to achieve more accurate large-scale models.

6

Chapter 6

Fluid systems

6

Chapter 6.1

Fluids in sedimentary basins: an overview R. Gaupp · P. Möller · V. Lüders · R. di Primio · R. Littke

6.1.1 Relevance of geofluids Sedimentary basins are compartments of the upper crust in which solid and fluid materials have accumulated over millions of years. Both, solids and fluids undergo partial transformation at temperatures ranging from 0° to 300 °C and pressures up to approximately 100 MPa. Due to their longevity and high contents of chemically metastable components, sedimentary basins can be regarded as longterm reactors, in which different fluids are transported, formed, mixed, modified, and - in the case of petroleum - consumed. The substance turnover and product composition of such a geo-reactor depends essentially on both externally and internally operating processes which affect the sedimentary basin fill over long geological periods. Fluids are transported by a variety of different processes, depending on pressure gradients and permeability distribution, but also on the presence of different types of fluids and their mixing. For example, migration of gas and oil through the pore space occurs either by pressure- or buoyancy-driven single-phase or multi-phase flow or by diffusion, being dependent on compaction, overpressure, fault zones and the geothermal field as well as associated fluid-rock interactions. Whereas groundwater flow in the shallow portions of sedimentary basins is generally dominated by differences in the topographic elevation of the water table/land surface (Galloway 1984; Hanor et al. 2004) with flow occurring from areas of higher elevation to lower elevation, fluid flow in deeper parts of sedimentary basins is greatly influenced by pressure gradients due to basin internal processes. Especially in the brittle zone, fluid flow is focused along highly permeable layers, fractures and faults. Subsurface flow systems can be very localised or regional. For example, large scale fluid flow over hundreds of kilometres involving water, oil and natural gas have been reported from North America and West Siberia (Bethke et al. 1991; Garven 1995; Littke et al. 1999). Geofluids in sedimentary basins provide the bulk of our drinking water, geothermal energy, and almost all oil and natural gas resources. Precipitation of mineral phases from fluids leads to mineral formation in the pore space

of sedimentary rocks or on fissures, including precipitation of precious metals. On the other hand, fluids can also dissolve minerals creating new pore space at depth. There are also more complex solution-dissolution reactions: for example, hydrocarbon fluids react with sulfates forming new inorganic gases and carbonates at certain temperatures. Finally, porous and permeable rocks are increasingly used as sites for storage of large amounts of fluids such as oil, natural gas, pressurised air and carbon dioxide adding to the enormous economic and ecological importance of sedimentary basins. In this chapter, some basic definitions on geofluids are given and selected important geochemical methods are presented which are used to investigate aqueous and petroleum fluids in the subsurface. Particular reference is given to the chemistry of groundwaters and formation waters of the North German Basin (NGB). One of the objectives is to close the obvious gap between research and literature on generation, flow, and composition of water on one hand and petroleum on the other, both using different parameters and definitions.

6.1.2 Definitions In geosciences the term “water” (or “waters”) mostly refers to an aqueous solution rather than to the chemical compound H2O. A general and consistent classification of aqueous fluids in sedimentary systems does not exist. The criteria are either descriptive (oil-field brine, basinal brine, basement brine), or based on salinity (ground- (= fresh), brackish-, saline-) or the origin of water (meteoric-, formation-, connate-). Water:

(a) the chemical compound H 2O; (b) any aqueous surface to subsurface solution which can be characterised by salinity or chlorinity or by adding terms characterising its production zone.

Salinity:

Synonymous with total dissolved solids (TDS), reported in mg/L (at 25 °C) or mg/kg.

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Chlorinity:

Corresponds to the concentration of chloride (Cl-) ions in either mg/L or mg/kg.

Meteoric water: Denotes any water derived from rain, snow, ice melt, streams and lakes percolating in rocks and displacing interstitial water. Formation water:

Water present in pores and fractures of rocks under natural conditions without any genetic or age specification.

Fluid:

Physical state in which a substance continually deforms (flows) under an applied shear stress; the term comprises both liquids and gases. In the context of sedimentary basins, subsurface fluids (geofluids) are liquids (waters, oil) with dissolved constituents as well as natural gases; a supercritical fluid is any substance at a temperature and pressure above its thermodynamic critical point; here a differentiation between liquid and gas phase is no longer possible and no variation of either pressure or temperature will result in the coexistence of a gaseous and liquid phase separated by a phase boundary.

Brine:

Hot, highly saline waters; a term widely used for pore fluids in deep sedimentary basins, oil-field waters and geothermal mineralising fluids.

Oil-field brine:

Saline water associated with hydrocarbon exploration and production.

Basinal brine:

Saline waters in basins.

Basement brine: Saline waters in basement rocks stored either in fissures and fractures or as fluid inclusions in minerals. Diagenesis:

encompasses all chemical, physical, and biological changes that sediments undergo after initial deposition and burial that convert a sediment into a lithified rock. Important processes are mechanical compaction, cementation, dissolution, mineral alteration. Weathering comprises diagenetic near-surface processes of alteration in sedimentary rocks. Processes that alter sediments at temperatures and pressures sufficiently high to be called metamorphism are excluded.

Water-rock interaction:

Water-rock interaction: covers all aspects of surface and subsurface waters that react with ambient rocks. The processes result in alteration of water chemistry and rock composition.

Connate water: Any kind of water that is of the same age as the sediment or rock in which it is stored. Groundwater = Salinity <1000 mg/L. Subsurface water of the saturation zone being distinct freshwater from surface water. Brackish water: Salinity 1000 to 10,000 mg/L. Saline water:

Any water with salinities exceeding seawater, i.e., 3.5% or 35,000 mg/L.

Thermo-saline water:

Water at elevated temperatures that has interacted intensely with sediments and/or rocks.

Compaction water:

connate or evolved water deposited with the sediment and meteoric water buried below the zone of active meteoric circulation expelled from sedimentary layers as a result of loading or compressive tectonic stress.

6.1.3

Subsurface aqueous fluids

6.1.3.1 Introduction Fluids in sedimentary basins play a fundamental role as media for the transport and/or redistribution of mass and heat. Depending on the burial history, regional thermal events, etc. fluids in porous sediments of different stratigraphic units can be very different from each other with regard to their chemical composition. Most of the fluids in the deeper parts of sedimentary basins are highly saline brines that could have different origins. For example, saline fluids that are produced for geothermal technology from Mesozoic strata of the North German Basin are not necessarily connate waters that were trapped during sedimentation. Instead, they are most likely fluids of meteoric origin that interacted with salt domes in the subsurface. Studies of stable isotopic compositions of formation waters

Fluids in sedimentary basins: an overview

from other basins support these findings (e.g., Clayton et al. 1966; Kharaka and Thordsen 1992). There are a couple of parameters such as topographic relief, tectonics, diagenetic processes, gravity-driven flux, metamorphism, devolatilisation, etc. that may induce a flow of geofluids at depth and/or drive fluids to lose their parental chemical signature during water-rock interaction. Upward fluidflow driven by subsidence (Bjørlykke 1983b, 1994) or channelised fluid migration along fractures and faults (Sibson 1994; Stephenson et al. 1994; Muchez and Sintubin 1998) may lead to mixing of highly-saline sedimentary brines with connate and/or even meteoric waters from shallower groundwater systems resulting in dilution of the ascending brines and salinisation of groundwater. Formation waters may show differences in salinity, isotopic composition and/or element content on a local scale (of hundreds of meters to kilometers), when compared with waters within the same strata but from other areas of the basin. Therefore, caution should be paid when the geochemical information obtained from one regional part of the basin is used for the reconstruction of the fluid evolution in the entire sedimentary basin. Furthermore, the direct examination of sedimentary brines stored at greater depth, for example Carboniferous strata in the Central European Basin is often limited to localities where such fluids are produced along with gases from deep-seated reservoirs. Reliable chemical data of deep basinal formation waters are rare.

Figure 6.1.1. Na/Cl vs. Br/Cl diagram. The Na-Cl-Br diagram can be used to discriminate between different solute sources (Kessler et al. 1995). Brines derived from dissolution of halite likely show Na/ Cl and Br/Cl ratios higher than seawater whereas brines produced by seawater evaporation past halite saturation have Na/Cl and Br/Cl ratios lower than seawater. Fluids that do not lie on the seawater evaporation or halite dissolution line are the result of water-rock interaction, when Na in the fluid is exchanged for another cation

The following chapter provides insights to the classification and origin of brines in general with emphasis on the North German part of the Central European Basin System (CEBS; see Maystrenko et al. this volume), as well as their chemical and isotopic characteristics.

6.1.3.2 Types of fluids There are many ways to describe the chemical composition of groundwaters and brines. Analyses of a variety of saline waters from various localities produced from different stratigraphic units are given in table 6.1.1. The simplest way is to characterise waters by their most abundant cations followed by anions. Na-Ca-HCO3: low-salinity weathering solutions; geothermal fluids Na-Cl: dominant dissolution of halite Na-Ca-Cl: diagenetic waters Na-Mg-Cl: seawater and seawater brines; Ca-Na-Cl: diagenetic water in limestone environments Ca-K-Cl: formation waters in contact with igneous rocks Mg-Cl: dissolution of post halite evaporites; extreme evaporation brines Na-SO4-HCO3: geothermal water in metamorphic rocks Analysis of cation and anion contents of fluids provide powerful tools to characterise the origin of saline fluids, to evaluate the degree of water-rock interaction, and to decipher the fluid evolution of basinal brines (Bethke 1996; Banks et al. 2000; Lüders et al. 2005). The gross chemical compositions and the regional or depth trends of saline waters may be visualised by Piper, Durov, Schoeller, and/or spider diagrams (Möller et al. 2006). Compared to limited species considered in Piper and Durov diagrams, Schoeller and spider diagrams encompass unlimited numbers of dissolved species. Combinations of various ionic ratios have also been proposed to characterise waters. During fluid flow Cl and Br behave conservatively and are not significantly affected by water-rock interactions except where the dissolution or precipitation of evaporites such as halite, sylvite etc. are involved. The involvement of Cl evaporites is unambiguous and brines produced by halite dissolution likely have Na/Cl and Br/Cl ratios considerably higher than seawater (Fig 6.1.1). The Na-Cl-Br diagram furthermore provides information about waterrock interaction. All samples in figure 6.1.1 that plot to the left of seawater evaporation or halite dissolution line have lost Na due to exchange of Na for other cations. This could be Ca if interaction with plagioclase is involved or K if albitisation of K-feldspar has taken place (Banks et al. 2002).

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mg/L

2,074 65 7.6

Sacramento Valley, California; Grimes, USA

mg/L

2,691 117 6.9

San Joaquin Valley, California,Wheeler Ridge, USA

mg/L

3,486 102 5.08

Central Mississippi Salt Dome Basin, Mississippi; Reedy Creek, USA

Kloppmann et al. 2001

5.5

156 5,870

183,000 65

314,000 0.5 121,000 1,290 1,340 1,270 24

mg/L

6.8

Stueber et al. 1993

12.7 2.5

529

82,800 164

134,800 4.9 45,300 155 1,330 4,360 191 2

mg/L

Mg-NaK-Cl 0.24 106 0.02 0.90

Fontes and Matray 1993

52

4,200

243,950 5,160

354,600 60 38,400 9,550 55,500 1,580 35

mg/L

236

741

Paris Basin, Varangeville mine, France

18,600 0.32 6,830 35.5 72 182 14.3 6.4 34

44,300 1.95 7,450 135 27 5,500 187 12 32 2 21,450 80 46 2,210 50 0.11 46

320,000 35 61,700 990 3,050 46,600 1,920 60 34 2 198,000 2,020 17 206 64 <0.02 28 59 Kharaka and Thordsen 1992

Na-CaCl 0.84 1136 1.99 19.8

Illiois Basin, Hord N, USA

11,000 44 30 359 <0.5 0.07 31

Kharaka and Thordsen 1992

Na-Cl 1.02 6337 0.58 0.51

North German Basin, Gorleben, GoHy483

Kharaka and Thordsen 1992

Na-Ca-Cl 0.48 221 9.28 495

Na-Ca-Cl 0.96 563 1.54

Na-Ca-Cl 0.54 603 123.75 7.38

Table 6.1.1. Compilation of chemical analyses of variously composed deep waters

depth T pH TDS Li Na K Mg Ca Sr Ba NH3 F Cl Br I HCO3SO42H2S SiO2 B

Literature Type Na/Cl molar ratio Cl/Br molar ratio Ca/Mg molar ratio Q value

AngaraLena Basin, Balagankinskaya 5, Russia

Menderes Graben, Kizildere, Turkey

Haon well, Lake Tiberias, Israel

24

AngaraLena Basin, Tuturskaya 1, Russia

mg/L

1122-1130

190 6.5

1790-1840

4.5

6.2

24,955

mg/L

5,212 442 1,657 1,078

mg/L 542,000

11,200 21,230 16,100 166,500

mg/L

17,360 20,550 13,300 144,220

450

504 60

15,800 202

230 590

Möller et al. 2007a

Na-SO4HCO3 14.23

3.44 0.005

Na-Mg-Ca-Cl 0.51 176 0.40 5.7

Möller et al. 2004

1.5 113

599,000

1,050

373,200 8,080 3 1,554 1,048

2,097 4 1,042 118 0.3 2

337,510 7,210 6 571 0.09

Ca-K-NaMg-Cl 0.05 104 6.28 176

Pinneker 1966 Pinneker 1966 Ca-K-NaMg-Cl 0.08 105 6.59 770

Fluids in sedimentary basins: an overview Figure 6.1.2. Lithological column of the NGB modified after Scheck and Bayer (1999). Abbreviations for the deep aquifer complexes are explained in the text

Other cation ratios such as Na/Li, Na/K, Ca/Mg also reflect the degree of interaction of palaeofluids with rocks in reservoirs and/or along flow paths (see also Schöner et al. this volume). The Na/K ratios in brines are likely to be controlled by equilibrium reactions with alkali feldspar and/or muscovite. Ca/Mg molar ratios increase during the evolution of brines by various reactions of brines with host rocks. For example, seawater shows a Ca/Mg ratio of 8.5; brines vary between >12 and about 0. Na/Cl molar ratios characterise the source of dominant Na salts such as from weathering yielding NaHCO3 (Na/Cl>>1), dissolution of evaporites - mainly halite - (Na/Cl=1) and presence of compounds such as CaCl2

and MgCl2 (Na/Cl<<1), which indicate the final stage of diagenetic and alteration processes.

6.1.3.3 Present-day fluids in the NGB The majority of water analyses of deep waters from in the eastern part of the NGB were produced in the course of hydrocarbon prospecting in the former German Democratic Republic in the 1970’s. These analyses were taken from the data files of the Geological Survey of BerlinBrandenburg and of commercial companies. Although the sampling sites are not evenly distributed, 291 analyses cover a diversity of brines in the NGB, and the various

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Figure 6.1.3. Map of the NGB showing locations of recent sampling (Möller et al. 2007b)

deep aquifer complexes (Deep Aquifer Complex DAC in Fig. 6.1.2). This set of analyses is supplemented by new analyses of waters from production and observation wells (Fig. 6.1.3; Möller et al. 2007b). TDS vs. depth

centre, depth of occurrence, spatial relation to salt domes, and thickness of geological formations strongly vary. For instance, the Permian evaporites are present at any depth because of salt diapirism. The situation is even more complex due to the local presence of Mesozoic evaporitic formations.

The cross plots of total dissolved solids (TDS) vs. depth of samples from NGB display two groups of samples in figure 6.1.4 (Hoth et al. 2005; Tesmer et al. 2007). Group 1 comprises samples from the Early Tertiary to Late Triassic aquifer complex (DAC 1) but also includes a few samples from other strata. TDS increases up to 270 g/L at depths of 2300 m and significantly correlates with depth. The gradient steepens at depths below 800 m. In group 2, brines from Late and Early Permian (DAC 3 and DAC 4) and Early Triassic (DAC 2) strata show no correlation of TDS and depth. The mineralisation of these waters (about 330 g/L) is by far higher than in group 1. Waters from DAC 3 are present in both groups but display a steep gradient of TDS only above 1000 m (figure 6.1.4). The samples from DAC 3 and DAC 4 are rather uniform in TDS. The stratigraphy does not relate to depth. Depending on the position of wells in the basin, border or

Figure 6.1.4. Cross plots of TDS and depth from Tertiary to EarlyPermian formations in the NGB (after Tesmer et al. 2007). Explanations are given in the text

Fluids in sedimentary basins: an overview Figure 6.1.5. Cross plots of Na+ vs. C1-. Two data groups emerge. Data for the marine evaporation line are taken from Fontes and Matray (1993)

Element correlations Cross-plots of Na+ vs. Cl- prove the existence of two groups of saline waters (Fig. 6.1.5). In group 1, Na+ and Cl- are positively correlated, whereas in group 2 the Zechstein water (DAC 3) shows an inverse correlation. Samples from the Bunter (DAC 2) are present in both groups with high Na concentrations. Samples from the Rotliegend (DAC 4) scatter at intermediate Na concentrations. The majority of samples do not plot along the seawater evaporation line after Fontes and Matray (1993). Similarly, cross plots of Br, Ca2+, and Mg2+ vs. Cl- reveal the existence of two groups, which both do not follow the marine evaporation line (Fig. 6.1.6). In general, group 1 is dominated by Na+ and Cl- with values up to 3,300 meq/L and 3,700 meq/L, respectively, and Ca2+ concentrations in the range between 50 and 150 meq/L, and Mg2+ concentrations vary between 10 and 180 meq/L. In group 2, Na+ may reach 5,000 meq/L but is no longer the dominant cation. Mg2+ and Ca2+ concentrations vary between 24 and 9,700 meq/L and 26 and 3,400 meq/L,

respectively. In group 2, two samples are outstanding because Mg2+ dominates the cation composition with more than 90% and Ca2+ is very low, Br- is either low or high. The HCO3- concentration is negligibly low (<15 meq/L) and SO42- concentration rarely exceeds 150 meq/L. In the cross plots of Na+/Cl- equivalent ratio vs. Clthe various trends in group 2 shows different trends in the involved aquifers. They are explicable by mixing of the following end members using PHREQC (version 2) (Parkhurst and Appelo 1999) (Fig. 6.1.7): modern seawater and seawater at beginning halite precipitation (curve 1), and saturated halite brine and evaporation brines at the beginning of sylvite (curve 2), carnallite (curve 3) and bischofite (curve 4) crystallisation. The majority of samples from the upper three deep aquifer complexes plot between curves 2 and 3. Water from DAC 1 plot between curve 1 and the seawater evaporation line (Fig. 6.1.6). Samples from DAC 2 plot either on the curve 2 or left of it. The majority of DAC 3 samples plot between curves 2 and 3. All DAC 4 samples (Early Permian) plot along or left of the seawater evaporation curve.

Figure 6.1.6. Cross plots of Ca2+, Mg2+ and Br- vs. Cl-. Note that the data do not plot on the seawater evaporation line after Fontes and Matray (1993), Tesmer et al. (2007). In the Ca-Cl diagram, the marine evaporation line follows the X axis

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In figure 6.1.9, the earth alkaline to alkali equivalent ratio (Ca2++Mg2+)/(Na++K+) increases over-proportional with Cl-excess, which demonstrates the continuous increase of Ca2++Mg2+ at the expense of Na++K+. The equivalent ratios of group 2 exceed by far those of group 1. Clexcess = [Cl-]-[Na+]-[K+] The Br-/Cl- equivalent ratios of both groups of waters plot in separate arrays, although they start at about the same values near Na+/Cl- equivalent ratios of about 1, i.e., dissolution of halite (Fig. 6.1.10). For orientation, the seawater evaporation line is given. The majority of samples of group 1 display Br -/Cl- ratios below the seawater ratio of 1.5. Group 1 is terminated when reaching the seawaters ratio of Br -/Cl-, whereas group 2 develops towards higher Br-/Cl- and lower Na+/Cl- ratios but the majority of samples plot below the seawater evaporation line (Fontes and Matray 1993). The spread in Na+/Cland Br -/Cl- equivalent ratios in group 2 is significantly larger than in group 1. The distribution of samples in group 2 shows distinct distribution with reference to their aquifers of origin.

Figure 6.1.8. Cross plots of Ca2+ vs. Mg2+ (Tesmer et al. 2007)

Spider patterns Four types of spider patterns are discernable in the new water samples from NGB (Fig. 6.1.11). Type S1: is characterised by dissolution of evaporite bodies: mainly dissolution of halite (Na/Cl >0.9), which yields Nasw≈ClswHCO3-sw, and Casw<Srsw indicate dominant dissolution of anhydrite and probable

Figure 6.1.9. Cross plots of (Ca2++Mg2+)/Na++K+) equivalent ratios vs. Clexcess (Tesmer et al. 2007). Clexcess represents the equivalents/L of Cl- not being balanced by dissolved NaCl and KCl. Clexcess quantifies the amount of dissolved CaCl2 and MgCl2, which can be expressed as mass of dissolved CaCl2 equivalents

Fluids in sedimentary basins: an overview

Figure 6.1.10. Cross plots of Br-/Cl- equivalent ratios vs. Na+/Clequivalent ratios (Tesmer et al. 2007)

mixing with Ca-Cl brine, and Bsw>SO42-sw≈Nasw points to strong sulfate reduction; Type S1b: Casw≤HCO3-sw, Casw≈Srsw or Casw>Srsw, and Bsw<SO42-sw, show much more smoothed patterns than in type S1a; Type S2: indicates dissolution of dominant carbonate with less dissolution of anhydrite and halite. This type is characterised by CaswSrsw. Type S3: The trend Ksw≈Nasw≈Clsw≈Srsw indicates marine formation water with little additional dissolved carbonate or infiltrating seawater in a coastal aquifer. Normalised values of 0.1 indicate 10 times dilution by fresh water. Type S4: with HCO3-swBrsw indicates dissolution of halite and sylvinite along with minor anhydrite and very little carbonate. The S-type waters are not specific for geological formations, but reflect various degrees of dissolution of carbonate, sulfate, and chloride minerals, largely independent of the present lithology. REY distribution patterns C1-chondrite-normalised rare earth element and Yttrium (REY) patterns are grouped according to their shapes (Fig. 6.1.12). Three different types emerge: Type R1: with concave-shaped patterns are controlled by oxygen-rich waters due to which FeOOH is precipitated during weathering of Fe(II)bearing minerals (Bau 1999; Kawabe et al. 1999; Quinn et al. 2004). In any case, these recent, oxygen-rich waters represent interaction with Quaternary sediments. Type R2: decrease from La to Eu and flatten thereafter. This type of pattern is typical for middle to late diagenetic limestones, geologically spoken Cretaceous and older limestones (Möller

Figure 6.1.11. Examples of spider patterns: type S1a: Neubrandenburg/#3; type S1b: Gorleben; type S2: Berlin, Köpenicker St./#1; type S3: Dummersdorf Ufer/#23; type S4: Lüneburg/#31 (Möller et al. 2007b)

Figure 6.1.12. Examples of rare earth element distribution patterns. Type R1: Travenbrück/#25; type R2: Berlin/Reichstag/#11; type R3: Rheinsberg/#32 after Möller et al. (2007b)

Type R3:

et al. 2003). Water-rock interaction with limestone involves dissolution of calcite. Therefore, water and limestones show the same type of pattern. Type R2 is typical for waters derived from Cretaceous to Triassic limestones in the study area. patterns with high light Rare Earth Elements (REE) content that decreases from middle to heavy REY (Rare Earth Elements and Yttrium; Fig. 6.1.12) and are typical for waters that experienced enhanced temperatures in water-rock interaction, i.e., for waters from depths ≥1000 m (Möller et al. 2003, 2006).

Like for S types, R types are not related to specific geological formations and their rock compositions. Water isotopes In figure 6.1.13, most of δ2H and δ18O of modern groundwater plot along the global meteoric water line (GMWL; Craig (1961)). Mean isotope composition of modern groundwater in northern Germany shows δ2H and δ18O

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Box 6.1.1 Distribution patterns of rare earth elements Conventionally, the abundance of rare earth elements (REE) is presented in semi-logarithmic plots of normalised concentrations vs. the atomic number or ionic size of REE on the X-axis. Rare earth elements together with yttrium are henceforth termed REY. C1-chondrite, Post-Archean Australian shale (PAAS), North American shale composite (NASC) (Anders and Grevesse 1989) are commonly used for normalisation. The average of carbonaceous C1 chondrite shows the least fractionated REY composition. In this contribution, normalisation to C1 chondrite is preferred because its REY abundance is the least fractionated known today (Fig. 1). In contrast, normalisation to shale composites such as PAAS and NASC alters Eu anomalies by increasing positive and decreasing negative ones because shales always show negative Eu anomalies with respect to C1 chondrite. Although not being a rare earth element (REE), yttrium is included in REE distribution patterns, because Y has the same ionic charge as Ho and nearly the ionic size of Ho. Y behaves like Ho in igneous rocks, but not so in chemical precipitates (Bau and Dulski 1995). In aqueous systems Y and Ho are slightly fractionated due to small differences in solution and surface complexation, and solubility products of the hydroxides (Diakonov et al. 1998). In waters, the distribution patterns of REY depends on soluble components or leachability of REY from less soluble minerals in rocks and sediments and the adsorption of REY by mineral surfaces along the pathway of water and/or in aquifers. In general, the REY patterns of water do not mirror those of the rocks or sediments. Their shapes are mostly significantly different. The only exception is water released from limestone and chalk because their dominant component calcite is dissolved and REY are released unfractionated. However, all waters show much lower REY concentrations than their parental limestones. By water-rock interaction the most soluble REY-bearing major and accessory minerals are the sources of REY but their concentration in water is strongly controlled by adsorption onto mineral surfaces (Möller et al. 2003). In REY patterns many elements may show anomalous behaviour. Ce(III) can be oxidised to Ce(IV), Eu(III) can be reduced to Eu(II). Besides these redox-controlled changes of concentrations, other anomalies are due to fractionation of REY along with chemical reactions in the hydrosphere. For instance, REY are adsorbed onto fresh precipitates of FeOOH and MnO2 under oxidising conditions. Strongly depending on pH, tetrad effects may be produced as a result of which La, Gd, Y and Lu are relatively enriched in water (Bau 1999; Kawabe et al. 1999; Quinn et al. 2004). This process plays a role in weathering of Fe and Mn rich rocks such as basalts. To some extent this also happens during karstification of Fe-rich limestone under oxidising conditions. Under reducing condition, this type of fractionation does not occur. Besides anomalies created by the above reactions, anomalies can also be inherited from the weathering of rocks. For example, groundwater originating from granites usually is characterised by a negative Eu anomaly, which results from weathering of biotite and many accessory components with strongly negative Eu anomalies (Möller et al. 1997). Metasedimentary rocks are often enriched in Ce as part of Fe oxide coatings on minerals which is the result of weathering of the precursor sediments under oxidising conditions. In cases where these metasedimentary rocks or their clastic debris experience reducing condition (which is common at increasing burial of sediments), Ce is released and waters acquire a positive Ce anomaly (Fig. 2).

Figure 1. Rare earth distribution pattern of shale normalised to C1 chondrite. Note the negative Eu anomaly which is gained by clastic sediments due to preferential leaching of Eu

Figure 2. Lanthanide anomalies in saline waters. All patterns show variable Y anomalies. The saline waters from Tabigha (Lake Tiberias, Israel) originating from Eocene limestones show strong negative Ce anomalies largely inherited from seawater, whereas the Mezar water from Cretaceous limestones (Golan, Israel) shows none (Möller et al. 2003). Groundwater from the Rotliegend conglomerates only show small anomalies, whereas water from the Rotliegend volcanic rocks shows strong positive Eu, Gd and Y anomalies

Fluids in sedimentary basins: an overview

The thermo-mineral waters (type R3) represent the interaction of early melt water with pre-Cretaceous limestones and mixing with their formation waters. More recent precipitation leached Sr from K-feldspar and these shallow waters of type R1, which locally interacted with marls and limestones during infiltration resulting in type R2. These shallow groundwaters are salinised by relics of brines and/or by mixing with brines locally ascending from deeper sources.

Figure 6.1.13. δ2H vs δ18O diagram showing the different trends of data, grouped according to R type of waters after Tesmer et al. (2007). Gorleben samples are taken from Kloppmann et al. (2001)

values of about -60 ‰ and -8 ‰, respectively. Compared with modern precipitation (in Berlin: δ2H: -55.0 ‰; δ18O: -7.70 ‰) most waters indicate lower average temperatures than today. Any mixtures of recent precipitation (meteoric water) and Pleistocene precipitation or melt water from the ice shield overlays the global meteoric water line. The isotopically heavy deep brines (type R3) plot on a mixing line of meteoric and seawater represented by SMOW. This would indicate that the water of these mixtures is of pre-Pleistocene age. The shift of two samples to higher δ18O values may be due to either hydrothermal isotope exchange (Giggenbach 1992; Clark and Fritz 1997) or mixing with extreme marine evaporation brines (Sofer and Gat 1972). See also figure 6.1.8.

Sr isotopes yield insight into cycling of Ca as one of the most important elements in saline solutions. At closer look, the samples of type R1 and R2 in figure 6.1.14 can be split into three entities each of which is characterised by a typical range of Sr isotope ratio in limestones: (i) ratios of about 0.7094 assemble water from depths less

Figure 6.1.14: Sr isotopes vs. 1/Sr. Signatures of water samples refer to R types (Möller et al. 2007b). Gorleben samples are taken from Kloppmann et al. (2001)

Sr isotope ratio vs. 1/Sr Clastic debris of weathered rocks from the Scandinavian shield formed Quaternary and Tertiary sediments up to several 100 m thick in the NGB. Today they consist mostly of quartz, little K-feldspar and clay minerals. The biotite originally being present might have degraded to hydrobiotite and clay minerals with time. Thus, two important Sr sources were present: Biotite with both high Rb and Sr, and K-feldspar with low Sr but enough Rb in order to increase 87Sr/86Sr ratios by radioactive decay of 87Rb. Deglaciation might have enhanced initial weathering of biotite relative to feldspar, but the rate of biotite weathering slowed down with time and most of the radiogenic Sr was lost to the melt water (Blum and Erel 1995). Weathering of the more stable K-feldspar yielded low Sr solutions and may represent a much later process. Thus, an early and a late weathering solution interacted with Permian to Cretaceous limestones and Tertiary marls.

than 100 m, i.e., interaction with youngest carbonates only; this group also includes averages of brackish and fresh water samples from Gorleben; (ii) ratios between 0.7086 and 0.7082 of mostly R2 type may be controlled by Tertiary marls, and (iii) ratios between 0.7080 and 0.7076 (fully in the range of Cretaceous to Late Permian limestones) with R1 and R2 types of waters which also includes an average of saline waters and brines from Gorleben. The latter two groups are typical for waters that interacted with limestones or carbonate-rich boulder clay. The different nearly horizontal trend lines may indicate interaction with limestones of variable Sr contents. Sulfur and oxygen isotopes in sulfate Some correlation between δ34S (SO4) and δ18O (SO4) exists in the range of low isotopic values (Fig. 6.1.15). For δ34S (SO4)>20‰, δ18O (SO4) is virtually constant.

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R. Gaupp · P. Möller · V. Lüders · R. di Primio · R. Littke Figure 6.1.15. Stable isotopes of sulfate dissolved in water. Data are taken from Möller et al. (2007b). The grey field represents Gorleben Zechstein brines (Berner et al. 2002). Fields of Late Permian and Triassic/Jurassic sulfates are taken from Claypool et al. (1980)

Only few brines plot near to the fields of Late Permian and Triassic/Jurassic sulfates (Claypool et al. 1980) and the average of Zechstein sulfates at Gorleben (Berner et al. 2002). Values of δ34S (SO4) exceeding 30‰ indicate bacterial reduction of dissolved sulfate due to which the heavier species remain in solution. Compared to type R1 and R2, samples of R3 type cluster. In figure 6.1.16a, R1 and R2 type of waters cover the whole field, whereas the deep brines of type R3 water cluster. The samples with δ18O (SO42-) <5‰ from shallow depth and low sulfate concentrations could represent atmospheric input (Berner et al. 2002). This source, however, cannot account for samples with more than 100 mg/L. With the rate of oxygen isotope exchange in the sulfatewater system being very slow at low temperatures, other processes have to be considered. These samples may indicate sulfate-water O-exchange along with enzymatic processes during sulfate reduction (Fritz et al. 1989; Toran and Harris 1989; Van Stempvoort and Krouse 1994). The upper limit of O exchange is a δ18O value of about 28‰, which roughly corresponds to sulfate-water O-isotope equilibrium at about 15 °C (Berner et al. 2002). The Late Permian to Middle Triassic anhydrite was and probably still is dissolved by at least two isotopically different waters (Fig. 6.1.15): mixtures of precipitation and seawater or its evaporations brines (type R3), which yield isotopically heavy water now residing at depths below 1000 m as brines, and isotopically light mixtures of Pleistocene and Recent precipitation occupying the aquifers at lower depths. In figure 6.1.16 the isotopic composition of sulfate is related to SO42-/Cl- wt ratio. In both δ34S (SO4) and δ18O (SO4) vs. SO42-/Cl- plots typical sources of water can be outlined (Berner et al. 2002). Figure 6.1.16a elucidates that few samples still show Late Permian δ34S (SO4) (like

Gorleben) but most of them plot between the array of bacterial sulfate reduction and the Late Permian salts representing mixing. Type R1 and R2 waters scatter widely. In figure 6.1.16b only few samples plot in the Zechstein field. Most of type R3 samples plot in the field of “salinised shallow groundwaters” (Berner et al. 2002). In summary, REY and isotopes of the deep waters from Northern Germany (Neustadt-Glewe, Waren, Neubrandenburg, Bad Wilsnack, Templin and Belzig) are of the replacement type. Only the extreme Mg rich brines may represent evaporated seawater. According to their water isotopic composition the deep brines represent mixtures of seawater and meteoric water. Hydrochemical studies of deep aquifer systems elsewhere suggest that multiple sources of salinity may exist in the NGB (e.g., Nishri et al. 1988; Kloppmann et al. 2001). The high salinity of the deep aquifer systems is explicable by descendent meteoric water dissolving salt bodies, connate water dissolving salt bodies during the stage of salt diapirism, and relics of Permian evaporation brines. In conclusion, hydrodynamic and hydrochemical behaviour of brines are highly variable within the different aquifer complexes of the NGB. Especially the intensity of salt diapirism lead to a complex geological environment with preferential flow paths and a strong interaction between the temperature distribution, salt concentrations and groundwater flow fields. Subsequent erosion and glaciations complicated matters. Different flow regimes and the existence of different brines resulted in water-rock interaction and mixing with meteoric waters, which is comparable to processes being observed in the North American and Canadian sedimentary basin systems (Carpenter 1978; Stueber et al. 1993; Knauth 1988; Bein and Dutton 1993; Connolly et al. 1990; Evans and Nunn 1989; Benn0ett and Hanor 1987; Mehta et al. 2000).

Fluids in sedimentary basins: an overview Figure 6.1.16. Plot of δ34S and δ18O in sulfate as a function of SO42-/Clwt ratio. Data are taken from Möller et al. (2007b). Arrays are reproduced from Berner et al. (2002)

6.1.4 Petroleum fluids Petroleum fluids are mixtures of organic molecules of different molar masses and structures. Their physical state ranges from gaseous to liquid, their specific density is usually less than that of water, and their viscosity ranges from very low, e.g., in condensates, to high in heavy oils and tars. Different classification schemes are used to categorise petroleum fluids according to their physical and chemical properties. Many of these schemes, such as those based on boiling-point ranges and density, originate from the refining industry. The heterogeneous nature of petroleum fluids is due to their origin from different types of source rocks and the processes they undergo during migration and accumulation. Petroleum source rocks are usually fine-grained rocks with high percentages of organic matter (organic carbon contents are mostly > 2 wt-%, sometimes less; Tissot and Welte 1984) that is termed kerogen. Kerogen is defined as insoluble (in organic solvents) sedimentary organic matter, which is

derived from plants and to a lesser extent from other biomass (bacteria, archaea, animals). The principal pathway of petroleum generation from different precursor materials is shown in figure 6.1.17. Kerogen derives from original biomass, consisting of different major groups of natural organic substances (lipids, proteins, carbohydrates, lignin, etc.). These natural substances are partly strongly degraded and restructured to form kerogen and are partly preserved in an almost non-degraded form. The question, whether the first or the second pathway is the more important one has been discussed intensely during the last decade (see e.g., de Leeuw 2007). During early stages of burial kerogen is subject to microbial degradation in sediments, but the intensity of microbial activity decreases with increasing temperature and finally ceases. In deep sedimentary rocks, nutrient supply is very restricted, and therefore microbial activity is limited (Wilhelms et al. 2001; Larter et al. 2003; Larter and di Primio 2005). Within a temperature interval between 10 °C up to 60-80 °C, methane

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is formed by microbial processes such as CO2 reduction and acetate fermentation (Schoell 1980). This gas which is generated at shallow depth in the earth is commonly called “bacterial methane”; a better term would be “microbial methane”, because bacteria and archaea contribute to gas generation. Kerogen in sedimentary rocks is commonly classified according to the van Krevelen diagram, in which hydrogen/carbon ratios are plotted versus oxygen/carbon atomic ratios (Fig. 6.1.18; Tissot and Welte 1984). Alternatively, plots of the hydrogen index (HI) versus the oxygen index (OI) from Rock-Eval pyrolysis are used (di Primio et al. this volume). Type I kerogen is hydrogenrich (H/C>1.5) and characterised by abundant long-chain aliphatic moieties. It is relatively rare and often derived from algal precursors deposited under oxygen-deficient conditions. Many type I kerogens occur in lake deposits, in which degradation by sulfate-reducing bacteria is less pronounced than in sea water (due to the lower availability of sulfate) and in which water circulation in bottom water is restricted. In type II kerogens, aliphatic structures are still abundant, but H/C ratios are lower than 1.5. This kerogen is more abundant than type I, often characteristic of marine petroleum source rocks, and often derived from algal biomass that is more degraded than that in type I kerogen. Type II kerogen can also derive from a physical mixture of type I (algal) and type III (land plant) kerogen. Type III kerogen is characterised by high oxygen/carbon ratios and hydrogen/carbon ratios lower than 1.0. Long aliphatic chains are usually rare in this kerogen type, whereas short chains or methyl groups are very common. Accordingly this kerogen has only a limited oil generation potential, but a high gas (methane) generation potential. During early diagenesis, mainly carbon dioxide, water,

and oxygen-bearing organic molecules evolve from this kerogen. Type III kerogen is common and coals usually fall into the field of type III kerogen. It should be noted, however, that some coals have H/C ratios higher than one and hydrogen index values as high as 300 to 350 mg HC/g rock. Such kerogens are often termed type II-III. Finally, type IV kerogen is very hydrogen depleted and has almost no petroleum generation potential. With increasing temperature and maturation, kerogen looses volatile products such as hydrocarbons, carbon dioxide and water and evolves towards a carbon-rich residue termed anthracite or finally graphite. At greater depth and temperature, usually above 100 °C, thermal dissociation (“cracking”) and disproportionation leads to the formation of oil and thermal gas (Fig. 6.1.17). The lower temperature limit can be even less than 100 °C, if thermally very labile kerogen is present in rocks. Such labile kerogen is formed by natural vulcanisation, when sulfur is fixed in kerogen during early diagenesis. This process is most relevant, where sulfate-rich waters (marine water, hypersaline lake water) are in contact with organic matter-rich sediments in the absence of iron. In this case, reduced sulfur is not predominantly fixed in iron sulfides (pyrite and marcasite), but in kerogen. Such sulfur-rich kerogen is very labile and can generate hydrocarbons at very low temperatures (Fig. 6.1.19: Type II-S, Tegelaar and Noble 1994; di Primio and Horsfield 1996; Killops and Killops 2005). In contrast, type I kerogen is thermally most stable and generates petroleum at high temperatures and within a small temperature interval. Composition of crude oil depends greatly on kerogen structure on the one hand and temperature on the other. The first petroleum generated is often a relatively heavy Figure 6.1.17. Interrelationships between biomass, kerogen, crude oil, pyrobitumen and natural gas (modified after Tissot and Welte 1984; de Leeuw 2007)

Fluids in sedimentary basins: an overview

Figure 6.1.18. Kerogen classification in the classic van Krevelen diagram (left) and the modified van Krevelen diagram (right; modified after Killops and Killops 2005; Tissot and Welte 1984)

and heterocompound-rich product, whereas the latest stages of oil generation are characterised by greater abundance of short-chain aliphatics and small aromatic molecules, leading to a light product, finally condensate. Oils are classified according to different parameters, most commonly based on density, but also based on viscosity, sulfur content, and chemical composition. Density classification in petroleum industry follows the system of the American Petroleum Institute (API), defining API gravity (°) according to the following formula:

API = (141.5/specific gravity)-131.5 (at 15.5 °C or 60 °F). Here specific gravity is expressed in g/cm³. The gravity of petroleum ranges usually between 10 and 70° API; values higher than 45° characterise condensates and values lower than 25° mark heavy oils. Thus crude oils fall into the range of 25 to 45°, with values between 35 and 45° characterising light oils. Specific gravities of crude oils are roughly between 0.8 and 0.9 t/m3; condensates are lighter and heavy oils are heavier.

Figure 6.1.19. Kerogen transformation at a constant heating rate of 3°C Ma-1 for different kerogen types (modified after Killops and Killops 2005; Tegelaar and Noble 1994); calculated vitrinite reflectance (EASY%Ro) after Sweeney and Burnham 1990

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Table 6.1.2. Classification of petroleum based on API gravity Petroleum fluid type

API Gravity

Heavy crude (heavy oil)

<25° Medium gravity crude

25°-35°

light crude

35°-45°

conventional oil

Condensate

Another classification is based on abundance of compound groups within the petroleum. Straight and branched n-alkanes (paraffins), cyclic aliphatics (naphthenes), and aromatic hydrocarbons as well as hetero-compounds are plotted in a triangular diagram (Fig. 6.1.20). Hetero-compounds are defined as organic molecules containing not only carbon and hydrogen, but also nitrogen, sulfur, or oxygen (NSO-compounds). As terrestrial-derived kerogens are usually more aromatic in nature than marine-derived kerogens, they do not plot in the paraffin-corner of this diagram. Aromatic asphaltic oils can derive from all kerogen types, because secondary alteration such as biodegradation can modify original oil composition in such a way, that they plot the AA field of figure 6.1.20. Biodegradation is discussed in more detail in di Primio et al. (this volume). From an engineering point of view petroleum types are also classified based on their phase behaviour and physical properties. The main physical properties used are API gravity and the gas to oil ratio (GOR). GOR is usually expressed in volume/volume units whereby Scf/STB (standard cubic feet per stock tank barrel) and Sm3/Sm3

>45°

(standard cubic meters) are the most common volume units used (conversion factor is about 5.5). In petroleum engineering five different petroleum fluid types are generally distinguished: black oil, light oil, retrograde gas, wet gas and dry gas. Typical phase diagrams for these fluid types are shown in the PVT Modelling Box (di Primio et al. this volume). Black oils are characterised by GORs of 350 Sm3/Sm3 (2000 Scf/STB) or less. API gravities are commonly below 45°. The phase envelope (di Primio et al. this volume) of black oil is characterised by low saturation pressures (bubble point pressures) and extends to high temperatures. The critical point is usually at much higher temperatures than those prevailing in reservoirs. Such fluids are also called low shrinkage oils. Light oils have GORs between 350 and 600 Sm3/Sm3 (2000 and 3300 Scf/STB) and API gravities above 35°. Their phase envelopes are characterised by elevated saturation pressures and a critical point at pressure and temperature conditions close to those of the reservoir. Light oils are also termed high shrinkage oils or volatile oils. Figure 6.1.20. Composition of crude oils based on distribution of paraffins, naphthenes, and NSO compounds (modified after Tissot and Welte 1984)

Fluids in sedimentary basins: an overview Figure 6.1.21. Composition of natural gases of microbial and thermal origin. δ13C (methane) is a measure of the 12C/13C ratio of methane (see Schoell 1980). The values become more negative with increasing 12C-content. Data plotted versus wetness (A: methane / (ethane + propane)) and δD-values (B: measure of deuterium/hydrogen ratios). Data are differentiated for Carboniferous (circles), Rotliegend (triagles) and Zechstein (sqares) gases from the CEBS, but not for different regions within the basin (for details see Lokhorst et al. 1998)

Retrograde gases represent petroleum fluids which occur in the gas phase within the reservoir but separate into gas and oil (condensate) at surface conditions. Retrograde gases are also called gas-condensates, condensates or retrograde gas-condensates. The lower GOR limit for retrograde gases is around 600 Sm3/Sm3 (3300 Scf/STB), the upper limit is ill-constrained, reaching values as high as 20,000 Sm3/Sm3 (110,000 Scf/STB) in some cases. Usually when GORs are above 9000 Sm3/Sm3 (50,000 Scf/ STB) the amount of evolving liquid during production is so small as to become negligible. Surface liquids produced from retrograde gases have API gravities between 40° and 60° API. The phase envelopes of retrograde gases extend to lower temperatures and higher pressures than those of black oils or light oils. The critical temperature of

the fluid lies below the reservoir temperature. Wet gases are characterised by phase envelopes lying entirely below reservoir temperature. Wet gas occurs as a gas phase under reservoir conditions and only very minor amounts of liquid are separate at surface conditions. Dry gas is primarily methane with some higher homologues. Both, reservoir and surface conditions are clearly outside of the phase envelope. Thermal gas generation usually occurs at temperatures even higher than those necessary for oil generation. Natural gas is not only methane, but other hydrocarbon gases (ethane, propane, butane), carbon dioxide, molecular nitrogen and hydrogen sulfide are also common. The first hydrocarbon gas generated from kerogen is relatively rich

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in ethane, propane, and butane and is commonly called “wet gas”. At higher temperatures, methane becomes predominant (“dry gas” stage). This change in gas composition is used in natural hydrocarbon gas classification. In addition, data on the carbon and hydrogen isotope composition of the individual gas components (methane, ethane, propane) are used for classification. Probably the most widely used classification scheme is a plot of gas dryness (methane/(ethane + propane)) versus the isotopic composition of methane in terms of 12C over 13C (expresses as δ13C Fig. 6.1.21A) or the plot of δ13C versus δ2H (Fig. 6.1.21B). Data for the CEBS are plotted in Fig. 6.1.21 (after Lokhorst et al. 1998) and reveal regional differences and differences between the different main reservoirs in the Carboniferous, Permian Rotliegend Formation, Permian Zechstein Formation, and Triassic, respectively. However, a thermogenic origin can be deduced for (almost) all major gas accumulations. Microbial gas is characterised by high methane over ethane plus propane ratios and a high abundance of the light carbon isotope 12C. Thermogenic gas is initially more enriched in ethane and propane while the methane is depleted in the light isotope 12C (“isotopically heavy”) as compared to microbial gas. With increasing maturity, kerogen generates more and more methane relative to other hydrocarbon gases, and the methane becomes further enriched in the 13C isotope. Thermogenic methane in sedimentary basins has two major precursors. One is kerogen, from which methyl groups are released. This methane which is derived directly from

kerogen is called primary methane. The other precursor is soluble organic matter (bitumen). Bitumen is mainly generated from kerogen and includes all oil, but also heavier, tarlike products. If this bitumen is thermally further matured, it will be cracked into gas and a solid residue, ultimately to methane and pyrobitumen. The latter is a carbon-rich residual organic substance (see Taylor et al. 1998 for details). Gas derived from bitumen cracking is commonly called secondary methane. In general, temperatures necessary for secondary methane generation are higher than those for primary gas generation (di Primio et al. this volume). Upon burial and maturation, kerogen is not completely converted into oil and gas. Nevertheless, the conversion rate is high (>50%) for all reactive kerogen types. An exception is the least reactive, almost inert kerogen type IV. Mass losses have been calculated, both for coals (Littke and Leythaeuser 1993) and for type II kerogen (Rullkötter et al. 1988), showing that the bulk of product generation takes place up to a maturity of 2% vitrinite reflectance. Hydrocarbons are easily expelled from clastic source (Fig. 6.1.22), and less easily from coals, where part of the stored bitumen is trapped until conversion to gas takes place at even higher temperatures. TOC loss is in the range of 50% for type II kerogen. In the CEBS, a variety of gas and oil reservoirs have been detected during the past 140 years, including tars and heavy oils, mildly biodegraded black oils, wet gases and dry gases with variable contents of inorganic gases, especially molecular nitrogen, hydrogen sulfide, and carbon dioxide. An overview on gas has been published by Lokhorst et al.

Fluids in sedimentary basins: an overview

(1998) and more information is found in di Primio et al. (this volume) and Krooss et al. (this volume). In general, accumulations of natural gas, including large accumulations of nitrogen-rich, carbon dioxide-rich and hydrogen sulfide enriched gases are found in the deeper and older part of the sedimentary succession, in particular in Late Triassic (Buntsandstein), Permian (Zechstein and Rotliegend) and Carboniferous reservoir rocks. A map showing the distribution of gas reservoirs is shown in figure 6.1.23 (after Lokhorst et al. 1998 and others). With the exception of the Zechstein carbonates, all reservoirs are predominantely clastic sandstones. Due to deep burial and/or cementation, permeabilities are often in the sub-milli Darcy range. These gas-filled tight reservoirs are currently intensely explored due to their economic potential. The gas in the reservoirs is to a great extent derived from Late Carboniferous (Westphalian), thick, coalbearing sequences (Littke et al. 1995), but there are also

contributions from Zechstein and pre-Westphalian source rocks. Most of the reservoirs are along an east-west trending fairway which is characterised by deep burial and petroleum generation in the Late Cretaceous and Cenozoic. Oil fields are mainly present in Jurassic and Cretaceous rocks. Reservoirs include limestones and chalks, especially in the Danish North Sea area, but also clastic sandstones, which are predominant in the German onshore part. Principal source rocks are Jurassic Shales, such as the Late Jurassic “hot unit” in the Danish North Sea and the Liassic Posidonia Shale in the onshore area. Contributions from other units exist as well, such as from Wealden (lowermost Cretaceaous) shales in the Lower Saxony Basin. In parts of the basin, there is also oil in the Zechstein, which is regarded as a self-charging system: Carbonate reservoir rocks in this evaporate sequence grade laterally and basinwards into carbonate source rocks, in which petroleum generation took place (see Magri et al. this volume).

Figure 6.1.23. Map showing distribution of gas reservoirs in the Central European Basin system (after Lokhorst et al. 1998; LBEG 2006; TNO-BNO 2007; Danish Energy Authority 2006; Schöneich 1986; Domzalski and Mazurek 2003; PIG 2005)

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Chapter 6.2

Transport processes J.L. Urai · G. Nover · C. Zwach · R. Ondrak · R. Schöner · B.M. Krooss

6.2.1 Introduction The quantitative formulation of transport processes of energy (e.g., heat) or mass generally involves relating the driving force and the resulting flux with an appropriately chosen coefficient of proportionality: Mass or energy flux = coefficient x driving force The most relevant transport processes in sedimentary systems are1: Heat flow (Fourier’s law) Pressure-driven fluid flow (Darcy’s law) Diffusive transport (Fick’s law)

· · ·

The term “flux” denotes the quantity (of mass or energy) transported per unit time perpendicularly across a unit cross section area. In addition, coupled transport processes such as thermo-diffusion or thermal convection may occur. Theoretically, these coupled transport processes can be described by the Onsager reciprocal relations, used in non-equilibrium thermodynamics (cf. Mitchell 1976; Clennell 1997). Practically, however, the identification and quantitative description of coupled processes in the complex evolution of sedimentary systems is very difficult. Sedimentation and compaction processes, representing a combination of solid and fluid mass transfer in a gravitational field, constitute the major mass transport in sedimentary basins on the geological time scale (see Box 6.2.1). Therefore, their quantitative description and reconstruction is one fundamental aspect in numerical basin modelling. Fluid transport in sedimentary systems may proceed either within the interconnected pore system (matrix transport) of sedimentary sequences (unconsolidated material or consolidated rocks), within vugular void spaces as in certain carbonate rocks, or along fault and fracture systems. The corresponding transport regimes differ substantially in terms of scale and transport efficiency. While pressure-driven fluid transport is the more efficient transport mechanism, diffusive transport of molecules and ions 1

is of great relevance in the redistribution of matter over short distances (microns up to hundreds of metres on the geological time scale) and in (dead-) volumes that are not accessible to pressure-driven flow. Diffusion is frequently the rate-determining process in mineral and diagenetic reactions. This chapter addresses fluid and heat transport processes in sedimentary systems and the current state of knowledge concerning their quantification. The focus is on transport in the pore and fracture systems, under the aspects of porous reservoir rocks and low-permeability seals.

6.2.2 Physical mechanisms and concepts 6.2.2.1 Overview The ability of sedimentary rocks to accommodate and transport fluids is directly linked to their geological and tectonic history. Geological conditions such as depositional environment control the primary mineralogical composition, grain size, grain size distribution and grain shapes. The texture and pore system of rocks are subsequently modified by compaction, cementation and diagenetic or metamorphic overprint so that the petrophysical rock properties such as porosity, permeability and tortuosity are affected by numerous processes acting simultaneously or sequentially. Theoretical and practical aspects of the petrophysical characterisation of rocks are discussed in numerous textbooks (e.g., Bear 1972; Freeze and Cherry 1979; Tiab and Donaldson 2004) and therefore only a short recapitulation of the most important terms and concepts will be given here. These are: porosity (Φ): a measure of the void space of porous media permeability (k): a measure of the ability of porous rocks to transport fluids tortuosity (τ): a measure of the structure of interconnected pore channels and interstices.

· · ·

The endogenic and exogenic mass transport processes, such as erosion and deposition are not considered in this list.

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Box 6.2.1. Mechanical and physicochemical compaction Compaction of sediments is the reduction in volume of a sedimentary body that is caused by stress. This stress commonly results from the increasing overburden load of deposits accumulated during subsidence of a sedimentary basin. Compressional stress owing to tectonic movements may also cause compaction. Compaction involves a combination of mechanical and physicochemical processes and results in the decrease of initial porosity, permeability, and stratigraphic thickness, and the increase in density. Overburden stress is transmitted by grain-to-grain stress in uncemented sediments. The mechanical response to this stress includes the rearrangement/rotation of components, the plastic deformation of mechanically unstable particles, and the brittle failure of grains. Physicochemical compaction includes dissolution of components at points of contact and mineral transformations that result in a net reduction of bulk volume. The loss of porosity due to compaction is accompanied by expulsion of pore fluid. If the pore water pressure is equal to the hydrostatic pressure at any depth, and the sedimentary column is completely carried by the grain framework, the sediments are normally consolidated. If low permeability prevents the escape of pore fluid, compaction is retarded and the sediments are underconsolidated. Parts of the overburden load are then carried by the pore water, which builds up excess pore water pressure. Overconsolidation results from erosion of parts of the overburden, thus the sediments reflect a state of consolidation that is higher than expected from the remaining sedimentary column. Mechanical compaction in siliciclastic sediments is generally most effective in the uppermost 1-1.5 km of burial (Fig. 1). Clay-rich sediments have higher initial porosities than sand-rich sediments, but the latter are much less affected by mechanical compaction (Fig. 1). Sandstones rich in rigid grains such as quartz become less strongly compacted than sandstones rich in mechanically unstable lithic grains such as (meta-)pelites or volcanic clasts. Physicochemical compaction in clays involves dehydration of clay minerals, notably the progressive transformation of smectite to illite. Physicochemical compaction in sandstones takes place by dissolution of components at grain-grain contacts (“pressure solution”). Pressure-related dissolution of quartz is most effective at burial depth greater than about 2 km, but is also controlled by temperature, mineralogy and rock fabric (e.g., Houseknecht 1988; Bjørkum 1996). Pressure solution of carbonate is important for compaction of carbonate-rich sands and limestones. The interaction of mechanical compaction, physicochemical compaction and cementation is difficult to predict and leads to a wide scatter in measured porosity-depth relationships for individual case studies (e.g., Rieke and Chilingarian 1974, Chilingarian and Wolf 1975). To approximate compaction in sediments, Athy (1930) proposed the negative exponential equation: Φz = Φ0 · e-cz

(6.2.1)

where Φz is the porosity at the depth z, Φ0 is the initial porosity, z is the burial depth, and c is an empirical coefficient depending on the type of sediment. As a result of compaction, sandstones typically have higher porosities than shales at elevated burial depths. However, the reliability of porosity-depth curves is very limited if diagenetic processes and the heterogeneities in terms of mineralogical composition, grain size and sorting are not adequately considered. For shales, refined models have been developed to better account for clay mineral dehydration and permeability evolution (e.g., Broichhausen et al. 2005). Sandstone porosity may be strongly affected by cementation and/or dissolution. Early cementation, for example, strongly reduces porosity, but effectively hampers further compaction. Sandstone compaction can be more precisely described by intergranular volume, i.e., porosity plus volume of cements and matrix present in between the framework grains (Fig. 1). Intergranular volume vs. depth curves can help to predict compaction and porosity evolution for sandstones (Lander and Walderhaug 1999; Paxton et al. 2002). Figure 1. Schematic intergranular volume-depth relationships of sandstones and porosity-depth relationships of shales caused by mechanical and physicochemical compaction during burial. a) Example for expected percentage of mechanical compaction of a well sorted, medium-grained, quartz-rich sandstone (A); reduction of intergranular volume by physicochemical compaction (B); resulting intergranular volume-depth relationship (C); grey envelope represents the range of intergranular volume-depth relationships of various sandstones. b) Example for expected percentage of mechanical compaction of a shale (A); reduction of porosity by dehydration of clay minerals (B); resulting porosity-depth relationship (C); grey envelope represents typical porosity-depth relationships of shales. In nature, clay compaction is probably rather an episodic than a continuous process (e.g., Cartwright 1994)

Transport processes

6.2.2.2 Porosity

6.2.2.3 Permeability

Porosity is the ratio of the pore volume of a rock Vp to its bulk volume Vb:

If an interconnected pore or fracture system exists, fluids can pass through porous rocks under the influence of a pressure gradient. This property is denoted as permeability or, in hydrogeology, hydraulic conductivity. Permeability is a characteristic property of the porous medium and is determined by relating the volumetric fluid flow rate q (m³/s) through a given cross section area Ac (m³) to the pressure gradient, dp/dx (Pa/m) acting on the fluid. This relationship, commonly known as Darcy’s law, is written as:



(6.2.2)



(6.2.3)

where Vgr is the grain volume. Porosity is frequently reported as a percentage rather than as a volume fraction. Unconsolidated clay-rich (pelitic) sediments may have porosities up to 70% shortly after deposition while porosity values of gravels and sands are usually below 50%. Upon compaction and consolidation during burial and diagenesis the porosity is successively reduced and may approach values close to zero in deep shales and crystalline rocks. Typical porosity ranges for unconsolidated sediments and rocks are listed in table 6.2.1. During compaction the pore volume of sediments may be altered due to fluid-solid interactions (see Schöner et al. this volume): primary porosity refers to the porosity retained during burial and affected only by physical compaction; secondary porosity is additional porosity created by dissolution processes or fracture formation. Due to sedimentation and diagenetic processes, initially open and interconnected pores may become isolated. Thus, two further categories, the total (absolute) and the effective porosity have to be discriminated: the absolute porosity is the ratio of the total void space in the sample to the bulk volume, regardless of the degree of interconnection of the pores. It encompasses the volume of fluid-filled isolated pores, adsorbed water on grains or particle surfaces, e.g., clay. Effective (accessible) porosity thus denotes the fraction of the pore system that is interconnected and amenable to fluid exchange. This porosity may be further subdivided into “transport porosity”, the network of interconnected pores through which the volume flow of fluid phases occurs, and “dead-end” pores with stagnant fluid where exchange is possible only by diffusion (Fig. 6.2.1).

(6.2.4) where v (m/s) is the volume flux or “Darcy velocity”, μ (Pa s) the viscosity of the fluid and k (m²) is the permeability coefficient. Quite frequently permeability coefficients are reported in the unit “Darcy” rather than in the SI unit m². One Darcy (1D) is equal to 0.9869 µm² (0.9869·10-12 m²). Because permeability coefficients of sedimentary rocks extend over a range of more than 10 orders of magnitude, the terms millidarcy mD (~10-15 m²), microdarcy µD (~10-18 m²) or nanodarcy nD (~10-21 m²) are used. Table 6.2.2 lists the ranges of permeability coefficients of unconsolidated sediments and sedimentary rocks. The above definition of permeability (absolute permeability) assumes complete saturation of the pore system with one single fluid. If more than one fluid phase is present in the pore system, e.g., in oil/water and gas/water systems, the transport rates of the individual phases depend on their relative volume fractions (saturation) in the pore systems and are described in terms of relative permeabilities. The sum of the individual phase permeability coefficients, the effective permeability coefficient, is always less than the (single phase) absolute permeability coefficient of the system. Factors affecting permeability The ability of porous rocks to conduct fluids depends on porosity and the degree of interconnection of the pores. Layering and rock texture are controlled by depositional and diagenetic conditions and usually result in a smaller

Table 6.2.1. Range of porosity values for unconsolidated deposits and rocks. Porosity is expressed as fraction Unconsolidated deposits

Rocks

Gravel

Sand

Silt

Clay

Fractured basalt

Karst limestone

Sandstone

Dolomite

Shale

Crystalline rocks

0.4 0.25

0.5 0.25

0.5 0.25

0.7 0.4

0.5 - 0.05

0.5 - 0.05

0.3 - 0.05

0.2 - 0.0

0.1 - 0.0

0.1- 0.0

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or larger degree of permeability anisotropy (Fig. 6.2.2). Furthermore, permeability exhibits significant stress dependence as a result of pore volume decrease and fracture closing when rocks are exposed to high effective stress. The interaction of aqueous fluids with clay (swelling of smectites) may considerably reduce conduits or may even block transportation. Thus, permeability also depends on the chemical composition of the flow medium (brine). Therefore, laboratory based permeability measurements must consider the environmental conditions of the samples to provide reliable results of the in-situ permeability.

6.2.2.4 Permeability-porosity relationships Because permeability depends on the connectivity of the pore system rather than on absolute porosity, a general relationship between porosity and permeability does not exist. However, semi-empirical and empirical correlations can often be established for individual formations or lithologic units. In a semi-logarithmic plot of permeability versus porosity one often finds “straight” lines with different slopes for different rock types, e.g., coarse-grained to clay-rich sandstones or various carbonates. These differences in slopes reflect differences in grain size and grain distribution, fabric elements, texture and cementation. More sophisticated models take into consideration additional parameters such as grain-size distribution, specific pore surface area and pore dimensions (Table 6.2.3). Grain-based models correlate permeability with the square of grain size. Grain sorting may be used as an additional

parameter. Surface area models indicate that permeability is proportional to the inverse square of specific pore surface area, and pore-dimension models show permeability proportional to the inverse square of a pore dimension. The most popular and fundamental porosity-permeability correlation model is the Kozeny-Carman equation (Wyllie and Gregory 1955 and references cited therein). It expresses permeability as a function of two basic properties: porosity and specific surface area. In this model, the conduits are assumed to be capillary bundles. For a single cylindrical capillary of radius r the fluid flow rate is given by Poiseuille’s law:

(6.2.5)

The fluid flow rate (q) or flux (v) in a capillary bundle can be calculated by combining the Darcy (Eq. 6.2.4) and Poiseuille (Eq. 6.2.5) equations. One then obtains an expression of the following type:

(6.2.6)

Here A is the cross sectional area of a micro-channel, α is a dimensionless constant taking into account the geometry of the pore channel. If the concept of hydraulic radius (= ratio of pore volume to solid-fluid interfacial area) is applied, then the cross sectional area A is replaced by tortuosity τ (τ = (Le/L)²) and specific surface area Spore of the pores, thus yielding an extended version of the KozenyCarman relation:

(6.2.7) Figure 6.2.1. Structures of porous rocks: a) wellsorted uniform granular material with microporous particles; b) poorly sorted graded granular material with small particles filling the large pores; c) partially closed discontinuity systems occurring in an intact porous rock; d) open discontinuity systems occurring in an intact rock due to mechanic fracturing. (after Freeze and Cherry 1979 and Meinzer 1923)

Transport processes

Table 6.2.2. Magnitude of permeability coefficients (in Darcy) of unconsolidated deposits and rocks Rocks

Unconsolidated deposits Gravel

Sand

Silt

Clay

Fractured basalt

Karst limestone

Sandstone

Dolomite

Shale

Crystalline rocks

105

103 -10

1-10‑3

104 -10‑5

10 - 10-2

104 - 10

10 - 10-5

10 - 10-4

10-5 - 10-8

10-4 - 10-9

Here KT is the effective zoning factor, depending on pore size and shape, grain size and shape, their distribution, tortuosity and other factors. KT is a characteristic parameter for a given stratigraphic flow unit (see below). Within one geological unit one may find flow conditions in the reservoir being laterally and vertically continuous, thus having similar permeability, porosity and rock fabric characteristics. Such a system can be assumed to be one flow unit. Due to sedimentation conditions different flow units may be found within a formation, e.g., as a function of clay content in sandstones. Different flow units can be hydraulically connected to other flow units. To map such units, field techniques as well as laboratory data can be used. Amaefule et al. (1993) proposed a method for characterising flow zones with similar hydraulic properties. The method uses a modified version of the Kozeny-Carman equation and the mean hydraulic radius concept. It is applied on scales from the pore level

through the inter-well level and involves three terms that are convenient for use with conventional core analysis data:

· reservoir quality index

(6.2.8)

· normalized porosity index

(6.2.9)

· flow zone indicator



(6.2.10)

where RQI is given in µm. The FZI includes geological attributes like texture and mineralogical composition as well as pore geometry. According to Tiab and Donaldson (2004), samples with the same FZI values have similar pore throat size, and thus constitute a flow unit. Doublelogarithmic plots of RQI vs. NPI result in straight lines exhibiting different slopes for different flow zones. Figure 6.2.2. Sketch of a porous rock. Anisotropy: High vertical permeability, low horizontal permeability. Open pores, closed pores, dead-end pores, inaccessible pore volume narrow pore entry

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Table 6.2.3. Permeability models, k is the permeability coefficient, D is the grain diameter, σ is the standard deviation of the grain diameter, Φ is the fractional porosity, m is a formation factor giving the consolidation for sand and sandstones, C is a sorting index for grain diameters, Swi is a sorting index, Rc is a characteristic radius and Rh is the hydraulic radius Model Sand pack

Author Krumbein and Monk (1943)

Equation k = 0.76 Dg² e-1.31σ

Grain-based

Berg (1970)

k = 80.8D²Φ5.1 e-1.385p

Grain-based

Van Baaren (1970)

k = 10D²d Φ 3.64+m C-3.64

Surface area

Timur (1968)

Surface area

Sen et al. (1990)

k = 0.136 Swi-2 Φ 4.4 k = 0.794 Tl2.15 Φ m+2.15

Pore size model

Kozeny-Carman in Carman (1956)

k = 400 Rh² Φ m

Pore size model

Winland in Pittman (1992)

k = 49.4 R35² Φ 1.47

Pore size model

Katz and Thompson (1986)

k = 17.9 Rc² Φ m

6.2.3 Fault seals and top seals 6.2.3.1 Overview Top seals and fault seals have a major effect on the distribution of fluid pressure and accumulation of hydrocarbons in sedimentary basins, but many details of the processes and their dynamics are poorly understood. There is a reasonable understanding of the basic physical processes of retention, but the complex geometries and many feedback processes make quantitative prediction difficult. A large majority of proven seals are evaporites, shales or organic rich rocks. They are laterally continuous, ductile and have high capillary entry pressures. Sealing performance is thought to depend on the ability to resist fracturing. A ranking by lithology has been proposed as Salt > Anhydrite > Organic-rich shales > Clay shales > Silty shales > Carbonate mudstones > Cherts (Downey 1984). It is generally accepted that laterally continuous, ductile and impermeable rocks are favorable seals, and that lithology is the most important parameter in con-

trolling seal quality. The capillary resistance of a few inches thick clay shale is sufficient to trap a large oil column. However, a thicker seal is beneficial because it provides more chance of lithological continuity and protects against faults which offset the cap rock by more than its own thickness. For a full understanding of top seals and fault seals one has to consider both the macro- and microscale aspects of the system. At the macroscale these are the geometry of the structure, the lateral continuity of the sealing lithology and the complicated plumbing system in layers juxtaposed by faults (Fig. 6.2.3). At the microscale there are a number of mechanisms of single- or multiphase fluid flow through matrix- or fracture porosity (Fig. 6.2.4).

6.2.3.2 Micro- and macroscale processes On the microscale, one can distinguish two hydrocarbon transport mechanisms through a seal: separate phase flow and diffusional transport (Fig. 6.2.4). Separate phase flow can occur in (i) relatively permeable seal lithologies, with Figure 6.2.3. Schematic diagram illustrating the different pathways for leakage of hydrocarbons from a faulted structure. Hydrocarbons can leak over the spill point (1), along the fault (2) or across the fault (3). The rate of mechanisms (2) and (3) is determined by the transport properties of the fault zone

Transport processes

Darcy flow When the top seal is not water-wet, or when its capillary entry pressure has been reached, the seal will leak hydrocarbons by flow through its matrix porosity. At sufficiently low permeabilities considerable hydrocarbon accumulations may be supported. The flux out of the trap is a function of hydrocarbon column length and in dynamic systems with simultaneous charge such a system is to some extent self-regulating. Important parameters here are the seal’s second phase saturation and relative permeability. Flow through fractures By their capillary resistance, tight shales can hold very large hydrocarbon columns. It is clear that hydrocarbon leakage can only occur by other mechanisms. Migration through fractures is a likely candidate. Here, the two main classes are shear and extensional fracturing. Diffusion through the pore water The solubility of methane and light hydrocarbons in water is of such magnitude that diffusion of gas through top seals may become an important mechanism of leakage through geologic time. In basins where generation and expulsion were active over a long geologic period, the pore water may become saturated everywhere. Under these conditions diffusional fluxes are controlled by the equilibrium concentration gradient in a hydrostatic column.

6.2.3.3 Application to fluid-rock systems

Figure 6.2.4. Micro scale mechanisms of seal leakage, together with schematic illustrations of the typical pressure regimes

intermediate capillary displacement pressures, or in (ii) tight but brittle lithologies with connected fracture systems. Capillary effects In a water-wet porous rock the movement of oil or gas is hindered by the requirement of a pressure difference to force the hydrocarbon-water interface through the pore throats in the rock. For a rock with a given capillary entry pressure the maximum hydrocarbon column held by a capillary seal can be calculated if the properties of the hydrocarbon - water system are known. Watts (1987) describes applications of this principle to many field situations.

In a fluid-filled porous rock, capillary displacement of a second, non-miscible phase cannot be adequately described using a single pore throat diameter. This is done instead by specifying the capillary pressure curve which relates capillary pressure to saturation. The capillary displacement pressure pd is defined as the fluid pressure which produces a significant saturation (about 10%) and corresponding second phase permeability. This concept is illustrated in figure 6.2.5. Analogous to the arguments given above, the maximum hydrocarbon column held by a capillary seal can be calculated when the relevant rock parameters and fluid parameters are known. This has been done in the past by a number of authors (Berg 1975; Schowalter 1979; Watts 1987; Ingram et al. 1997). In this case, the maximum hydrocarbon column hmax held by a specific seal at a given depth is given by:

(6.2.11)

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Box 6.2.2 Capillary seals In a water-wet porous rock the movement of oil or gas is hindered by the requirement of a pressure difference to force the hydrocarbon-water interface through pore throats (Schowalter 1979; Berg 1975). For a single cylindrical pore throat, the pressure difference is given by (6.2.12) where pc is capillary pressure (= pressure difference across the fluid-fluid interface; Pa); γ interfacial tension (N/m); θ wetting angle (degrees) and r pore throat radius (m). For a stationary fluid, this relationship can be used to calculate the hydrocarbon column held by a capillary seal (with a single pore size): (6.2.13) where h is hydrocarbon column length (m), g = 9.8 ms‑2 and Δρ is the density difference between water and hydrocarbon (kgm-3). A simple demonstration of this process is shown when a small, commercially available sieve (mesh diameter = 0.6 mm) is inverted and immersed in water (Fig. 1). From air bubbles released under its rim, an air column is accumulated, retained by capillary resistance. After a maximum column length of about 35 mm is reached, addition of more air results in leakage of part of the air column. After some gas has escaped the seal is closed again by imbibition. Using equation (6.2.13) with γ = 72x10‑3 Nm-1, Δρ = 999 kg m‑1, r = 0.3 mm, and assuming θ = 30 degrees, one calculates a column length of 42 mm. This is in reasonable agreement with observation. However, a few points should be noted regarding this simple system: (i) seal thickness has no effect on the maximum hydrocarbon column needed for capillary entry; (ii) this is a static calculation: rates of leakage and sealing capacity in a dynamic situation are determined by other parameters. Figure 1. Illustration of the physical principles of a simple capillary seal, by an inverted sieve retaining an air column

where pdHC is the capillary displacement pressure for a hydrocarbon-water system. Two points are worth noting here. Firstly, hmax to a first approximation is independent of overpressure in the water phase. Figure 6.2.6 shows pressure-depth plots around two seals assumed to have identical pdHC at different depths. Both seals hold the same hydrocarbon column by capillary resistance, but in the shallower one the pressure at top reservoir is very close to the minimum

principle stress and the risk of hydrofracturing is high. Secondly, this analysis assumes full pressure communication between the wetting pore fluid in the lower part of the seal and in the reservoir. When the seal is slightly overpressured with respect to the reservoir, a larger hydrocarbon column can be held by the same seal. More generally, hydrodynamic trapping (Hubbert 1953) in basins with regional groundwater flow patterns can alter the trapping conditions and shapes of hydrocarbon-water contacts significantly.

Transport processes Figure 6.2.5. Schematic diagram illustrating the principle of capillary entry and associated second phase permeability

Figure 6.2.6. An illustration of the effect of different pore pressure systems on retention. PVT properties of the hydrocarbons are assumed to be independent of depth, and the displacement pressure of both seals is the same. Structural closure is larger than hmax. (a): the two seals hold the same hydrocarbon column, (b): the lower seal holds hmax despite the overpressure, but the higher seal is breached by hydrofracturing. (c): the lower seal holds a longer hydrocarbon column then in (a) because of the downward flow out of the seal. The upper seal is inactive in this case due to lack of charge

Displacement pressure Rock properties enter the calculations via the in-situ hydrocarbon-water displacement pressure pdHC. This parameter is usually determined by conversion of data from high pressure mercury injection in air-dried samples taken from cores or larger drill cuttings (Fig. 6.2.7). Direct measurement using gas and water in samples under in-situ stress and temperature conditions is possible

but much more time-consuming than in mercury-air tests (Katz and Lee 1990). From the displacement pressure of the mercury injection test (pdHg) the hydrocarbon-water displacement pressure is calculated as

(6.2.14)

Here the usual assumption is that the seal is water-wet (θHC = 0). In the mercury-air system, γHg = 0.480 N/m and θHg = 39 degrees.

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Capillary pressures and permeability in permeable seals may be coupled using the Leverett J function J(Sw) = pc/γ (k/Φ)0.5

show examples of mercury-air data from the Buntsandstein and Early Tertiary of the Central European Basin Systems (CEBS).

(6.2.15).

Two-phase flow can than be simulated using numerical techniques developed in reservoir engineering. To get an impression of the range of possible values in pdHg, on can compare values from reservoir sandstones (pdHg < 0.1 MPa) to shales. As shown by Smith (1966), significant columns can be retained by rocks with pdHg > 1 MPa. On the other end of the scale, in tight shales pdHg can be in excess of 50 MPa: these rocks could hold such large columns that seal capacity is rarely reached and seal leakage, when present, is by other mechanisms (Fig. 6.2.8). Large-scale two phase flow Two phase flow at large scale is a very complex process. Figure 6.2.9 shows schematic illustrations of the interplay between capillary and geometric processes which can result in separation of the flow paths of oil and gas in the subsurface. Rock properties One of the major sources of uncertainty in capillary seal prediction lies in estimating the Mercury-air displacement pressure pdHg. A capillary seals database was presented by Hildenbrand et al. (2006) (Fig. 6.2.10), who derived an equation to calculate pdHg from data on porosity and clay content, allowing models of depth-depen­ dent seal capacity in basins. It should be noted however, that although figure 6.2.11 shows a reasonable correlation for the set of Tertiary clays studied, extrapolation to other clays is not possible. Figure 6.2.11 through 6.2.13

Gas density Gas density ρg at a given pressure and temperature may be approximated using the equation of state for a methaneethane (C1 - C2) mixture:

(6.1.16)

where pg is gas pressure, R is the gas constant (8.3144 Jmol-1K-1), T is temperature in K, Z is the compressibility factor and M = 16.04 XC1 + 30.07 XC2 (mean molecular weight) (6.2.17) with X being the molar fractions of the components such that XC1 + XC2 = 1

(6.2.18)

Oil density In-situ oil density is less easily calculated, due to the large variations in composition. If sufficient well data are available, oil properties can be calculated using reservoir engineering algorithms or from RFT data. Alternatively, for existing fields one can use PVT determinations from production reports. For a first impression, oil density at surface and GOR can be converted to in-situ density using the relationships presented in Schowalter (1979). ρo = 1000((-0.00022 + 0.000002*(ρsurface-10))* (GOR-2000)+0.515)

(6.2.19)

Figure 6.2.7. Schematic illustration of the technique of mercury injection. A porous sample is immersed in Mercury. Then the pressure of the liquid metal is raised and at each pressure the amount of mercury entering the pore space is measured

Transport processes

Box 6.2.3 Leak dynamics When the capillary displacement pressure of a water-wet top seal has been reached, it will leak hydrocarbons by flow through its interconnected porosity. The dynamics of two-phase flow through a water-wet capillary fault seal with a displacement pressure of 1.55 MPa were modeled using the one-dimensional reservoir simulator ONESIM (de Vries pers. comm. 2005). Rock properties are as shown, reservoir and seal have different permeabilities, porosities and capillary pressure curves, but have the same relative permeability and a common Leverett J curve. Figure 1 illustrates the model and some results. Figure 1a is an illustration of sealing conditions at a fault where the imposed pressure difference is smaller than the capillary displacement pressure of the seal. At time = 0 water saturation equals 1 everywhere, and oil is injected from the left at constant pressure. The model is filled with oil up to the fault zone, and equilibrium is reached. In figure 1b all parameters are the same as in the previous simulation, but the pressure difference across the fault is now higher than the displacement pressure. After the reservoir on the left is saturated with oil, an oil front moves through the seal. In steady state there is a relatively low oil saturation in the seal. The seal is leaking oil at a rate lower than expected from its bulk permeability only, owing to the low saturation.

Figure 1. ONESIM models of a fault which is a capillary seal at differential pressures below 1.55 MPa, but leaks at higher differential pressures. Pressure on the left is higher. At time = 0 the water saturation equals 1 everywhere, and oil is injected from the left. Both reservoir and seal are assumed to have the same relative permeability curves, and a common Leverett J curve. Oil and water viscosities are 10‑3 Pa s. Simulations are run over 3000 years

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J.L. Urai · G. Nover · C. Zwach · R. Ondrak · R. Schöner · B.M. Krooss Figure 6.2.8. Typical mercury-air injection curves for sandstones and mudstones

Figure 6.2.9. (a) Illustration of the large-scale effects of capillary seal leakage during gas (green) and oil (red) migration, showing how in a carrier bed with variable capillary entry pressure different oil and gas accumulations evolve over time; after Schowalter (1981). (b-e) Structural closure with a leakage path across a fault, showing how the interplay between rates of charge of oil and gas and the rates of leakage across the fault can produce widely different hydrocarbon accumulations. (green = gas, red = oil, yellow = water)

Transport processes Figure 6.2.10. Correlation between clay content, porosity and mercury-air Capillary entry pressure for a suite of Tertiary clays. After Hildenbrand et al. (2006)

In this empirical formula, ρo is in-situ oil density (kgm-3), ρsurface is the density of the oil at surface (API) and GOR the gas-oil ratio. The conversion from API units to kgm-3 is: ρo = 141500/(API+131.5). Interfacial tension Accurate theoretical prediction of interphase tensions in hydrocarbon-water systems is a difficult and as yet un-

solved problem. The two main parameters are density contrast and mutual solubility. Estimations are currently based on laboratory experiments but these are also problematic and sensitive to subtle changes in composition and interface structure. This has strong effect on the maximum columns calculated; interfacial tension is an important source of uncertainty in capillary seal calculations. Examples of interfacial tension data are presented in Schowalter (1979) and Watts (1987). Figure 6.2.11. Envelope of mercury-air capillary curves from silty mudstones, muddy siltstones and mudstones of the Buntsandstein in the CEBS. Samples are from the depth range 800 – 4000 m

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J.L. Urai · G. Nover · C. Zwach · R. Ondrak · R. Schöner · B.M. Krooss Figure 6.2.12. The same dataset as in figure 6.2.11, showing the capillary displacement pressure (capillary pressure at 10% saturation), grouped to different sedimentary environments. It can be seen that there is a large scatter in displacement pressures, so that sedimentary environment is a poor predictor of seal capacity in this setting

Wetting properties In parallel with the uncertainties in interfacial tensions, wetting is another poorly understood parameter in hydrocarbon-water-rock systems. Recent reviews are by Anderson (1986) and Swanson (1981). Note that wetting is a term related to contact angles in a solid-two fluid system, and does not directly relate to the fluid in contact with the rock at a given time. An important point relevant to capillary seals is that clean silicate surfaces are water-wet. Silicate rocks which have been exposed to hydrocarbons, and other minerals (carbonates, anhydrite) may be mixed-wet or oil wet. In silicates this change is associated with the adsorption of a variety of compounds from crude oil (such as asphaltenes, see Anderson 1986).

through the pore fluid) is the most likely leak mechanism. Especially in the case of gas, a few thin fractures will allow depletion of a large field over geologic time. For this reason, ductility and its opposite, brittleness, are seen as key properties of effective seals. In hydrocarbon seal studies, the terms brittle and ductile are used somewhat differently than in traditional rock mechanics. Because the key parameter here is the ability to transport fluids by one- or two phase flow, the following definitions are used (Ingram et al. 1997): a ductile shale is able to undergo plastic deformation without increasing its permeability (it may contain non-dilatant, impermeable fractures), whereas a brittle one will develop dilatant fractures and increase its permeability by many orders of magnitude when deformed (Fig. 6.2.14).

By their capillary resistance, many seals can hold large hydrocarbon columns. From these, hydrocarbon leakage can occur by alternative mechanisms. Migration through fractures (tension fractures or shear fractures) is a likely candidate.

A brittle seal does not necessarily leak hydrocarbons: in the absence of tectonic deformation after the trap is filled, even very brittle lithologies can form excellent seals. Alternatively, initially dilatant faults may become re-sealed by processes such as the formation of a clay gouge (Holland et al. 2006) or cementation by hydrothermal fluids (Nollet et al. 2005). Therefore, an evaluation of the risk of seal leakage along fractures must always involve a structural and charge analysis.

6.2.3.5 Brittle and Ductile Seals

Thus, a key to analysing shale top seals is to quantify the conditions where dilatant mechanical failure occurs.

Unfractured compacted mudrocks can hold very large hydrocarbon columns. In these rocks hydrocarbons can not move by Darcy flow through matrix porosity; flow through permeable fracture networks (besides diffusion

This requires quantification of two competing parameters: (i) mechanical properties of the formation under investigation, as given by its cohesion (C), friction angle (φ)

6.2.3.4 (Hydro) fractured seal

Transport processes

faults. The two end-member cases here are flow along the fault (fault is completely open) and no flow in any direction (fault is completely sealing). The case in-between where there may be a seal along the fault but flow across the fault, or even more complicated systems of limited sealing over shorter distances in either direction, are the most complex. On the timescales of hydrocarbon production, the relative importance of sealing and non-sealing faults is quite different, and faults which have not played a role in the accumulation of oil or gas in a structure may form important flow barriers. The first step in evaluating possible across-fault flow pathways consists of an analysis of possible juxtapositions across the fault. Although the principle can be illustrated in profile, the analysis is necessarily a 3D process, considering the topology of the fault system (Fig. 6.2.15) and the variation in offset along the fault. The final analysis is usually carried out using a fault plane diagram constructed in a plane parallel to the fault plane.

Figure 6.2.13. Mercury–air capillary curves from Tertiary claystones in the CEBS

and unconfined compressive strength (UCS), (a stronger rock is more brittle); (ii) the state of effective stress in the seal (high effective stress enhances ductility).

6.2.3.6 Fault Seals Many hydrocarbon accumulations are affected in some way by the presence of faults in the reservoir and/or seal lithologies. In exploration and appraisal workflow, evaluation of the different leak pathways from a structure usually involves an assessment of the transport properties of

In the case of juxtaposition of permeable lithologies (Knipe et al. 1997) the transport properties of the fault gouge are the critical parameter controlling fluid flow. Prediction of the internal structure of fault gouges is difficult, due to the large amount of highly complex processes involved. The initially segmented geometry of fault zones and subsequent coalescence of segments, in combination with the contrast in mechanical and transport properties between the layers are the two most important factors in this evolution. The resulting ranges of structures are illustrated in figure 6.2.16. It is clear from the above that fault zones can have vastly higher or lower permeabilities than their country rocks, and thus form either barriers or conduits for fluid flow. The amount and type of clay in a fault is one of the main parameters controlling the mechanical strength and fluid transport properties of the fault gouge (Logan and Rauenzahn 1987; Lupini et al. 1981; Heynekamp et al. 1999; Rawling et al. 2001). Much of the current re-

Figure 6.2.14. Illustration of the top seal above an oil accumulation, which becomes deformed after charge. In the first case the seal is ductile and can deform without becoming permeable: the accumulation is not affected. In the second case the deforming seal becomes permeable, and the accumulation moves to a higher structure

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Box 6.2.4 Seal failure The sequence of calculations and the equations used are given below. Further details are given in Ingram and Urai (1997): Unconfined compressive strength (UCS), MPa, from sonic: (6.2.20) where UCS is expressed in MPa and Δt is sonic transit time in microseconds per meter. Friction angle (φ), degrees, from surface area: log(φ)=1.544-0.00136S

(6.2.21)

where φ is friction angle (degrees) and S is the total surface area in m2/g. (Surface area is determined using the DCM method; highest values obtainable are for sodium montmorillonite at 800 m2/g). Cohesion (C), MPa from UCS and φ: (6.2.22) Vertical effective stress (σ'v), MPa: (6.2.23) where ρi is bulk density minus fluid density, g=9.81 ms-1 and ∆ z is the length of the interval considered. Assuming a Mohr-Coulomb failure criterion in an extensional setting, the minimum horizontal effective stress (σ'h min, MPa) at which the seal starts to deform plastically can be calculated as: (6.2.24) σ'3 and σ'1 are the minimum and maximum principle effective stresses (also defined in 6.2.3.6). For the case of σ'v > UCS, σ'h min is positive and all effective stresses at failure are compressive. Therefore, fracturing is mixed-mode or in shear. For the case of σ'v < UCS, σ‘h min is negative and fracturing is in extension and thus highly dilatant. In the extreme situation where the condition pf = σ3 + τ (6.2.25) is satisfied, hydraulic fractures will form and allow the hydrocarbons to escape. In this equation, pf is the fluid (hydrocarbon or water) pressure at the top seal-reservoir interface, σ3 the minimum principle (total) stress and τ the tensile strength (of the top seal).

search on fault seal is therefore focused on the process of clay being incorporated into the fault gouge, often called “clay smear”. Clay smear (Fig. 6.2.17) is a loosely defined term born in hydrocarbon geology (van der Zee 2002; Aydin and Eyal 2002); its usage differs between publications and the definition of processes operating is often unclear. In the most general meaning, the term includes all processes which somehow transform clay in the wall rock into clay which is part of the fault zone. Processes included are clay abrasion (Lindsay et al. 1993), shear in releasing fault links (Bashir Koledoye et al. 2000), preferred smear and lateral clay injection (Lehner and Pilaar 1997; van der Zee et al. 2003). For applied studies of clay smear evaluation a number of semi-empirical tools are available (Bouvier et al. 1989; Fulljames et al. 1997; Lindsay et al. 1993; Fristad et al. 1997). An overview and comparison of these methods is

given by Yielding et al. (1997). Most of these methods are based on the assumption that the fault gouge is a reworked equivalent of the wall rock without addition or removal of material, and therefore has, on average, the same clay fraction. Unfortunately this assumption is statistical in nature, and it does not specify which of the different structures shown in figure 6.2.16 is present. Therefore, predictions based on this method have a high degree of uncertainty because rock properties are not included in the calculation. A simplified model in Mohr space To illustrate the main mechanical aspects of clay injection, we present a simple model using Mohr diagrams (Fig. 6.2.18). The structural evolution is at a stage where the fault has just formed across the whole area considered, but the fault throw is still very small and the two sides of the

Transport processes

Figure 6.2.15. (a) Schematic diagram showing a reservoir sand containing two faults bound by a tip line (dashed line). The two faults are separated by a relay ramp. (b) Vertical profile across one of the faults in (a), now showing several stacked reservoir layers. Fault tip is shown by blue dots. Note the increase in fault throw from the fault tip to the centre. The hanging wall and footwall parts of the layers are shown in different shades. (c) Fault plane diagram drawn in a plane parallel to the fault plane, showing juxtaposition of footwall and hanging wall units. Tip line of the fault is shown by dashed line. In the case of across-fault flow several of the layers are brought into communication by the presence of the fault

fault in the releasing bend have been partially unloaded but have not yet separated. We now consider a point (A) in the releasing bend. The maximum principal effective stress (σ'1) approximately equals the overburden load. The minimum principle effective stress (σ'3) in this point is approximately horizontal, and is progressively reduced by the fault movement. In Mohr space this results in the Mohr circle with its right leg fixed and its left leg moving to the left. At this point there are two possibilities: (i) The yield envelope is reached before σ'3 = 0, and plastic flow of the clay towards the pull apart structure is initiated. (ii) In a more cohesive clay, the yield envelope is not reached before σ'3 = 0. The pull-apart structure will open because the clay layer can support the vertical load without becoming plastic. Therefore, there will be no injection, but the formation of a dilatant void instead. In this location other processes may take place such as vein growth (Hilgers et al. 2004). For the case of a linear Mohr-Coulomb plasticity criterion the condition for the onset of lateral injection can be written by:

C = σ'1 (1 - sinφ)/(2 cos φ)

(6.2.26)

Here C is cohesion (MPa), σ'1 is the maximum principle effective stress (MPa) and φ is internal friction angle (degrees). Equation 6.2.26 is a first-order description of the conditions required for injection of clay. It contains a description of the interdependence of in-situ stress and rock strength. The usefulness of this equation comes from the fact that all the parameters involved can be in principle obtained from wire line log and cuttings data. Therefore, this relationship can be used to predict the onset of clay injection in the subsurface.

6.2.3.7 Numerical modelling of petroleum flow The flow of petroleum in sedimentary systems is dominantly governed by the interplay of forces which on the one hand drive and on the other resist movement. Petroleum generated in mature source rocks migrates within these usually fine-grained rocks until it reaches more

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Figure 6.2.16. An illustration of the main structural elements found in fault zones in sediments (after van der Zee et al. 2003). (a) heterogeneous simple shear deformation, without strong localisation in a discrete fault plane; (b) heterogeneous simple shear, with lateral transition into a sharply localised zone of deformation; (c) lens structure between two branches of a fault plane, usually showing a higher degree of deformation than the surrounding rocks; (d) fault gouge consisting of parallel strands, with telescoping of the individual layers; (e) fault gouge with lateral injection of a weak layer (this paper); (f) fault gouge formed by the process of preferred smear of the soft layer; (g) fault gouge formed by disruption of the stronger layer followed by mixing of the fragments; (h) the brittle end member fault, with open fractures, developing preferentially in releasing sections

porous and permeable carrier rocks. These two closely linked processes are commonly referred to as “primary migration” and “expulsion” (Welte et al. 2000). Petroleum flow within the permeable carrier rocks until a trap is reached is referred to as “secondary migration”. Longand short-term leakage processes out of reservoir structures through the seal (capillary leakage), along fault and fracture systems or past “spill points” are referred to as “tertiary migration” (Sylta 2004).

Modern basin simulation software packages use one or more of the four following concepts to model fluid flow under subsurface conditions on the geological time scale:

The main driving force for migration is the buoyancy of the petroleum phase, as a result of its lower specific density compared to surrounding pore water under subsurface conditions. Overpressure gradients in the continuous pore water phase can add to this flow drive (Hubbert 1953). Viscous and capillary forces resist this driving force. The magnitude of the viscous forces is controlled by the dynamic viscosity of the petroleum phase and the relative permeability of the rock matrix with respect to this phase. The capillary forces manifest themselves in terms of the capillary entry pressure which depends on the wettability properties of the pore system and interfacial tension between the fluid phases. Fluid flow will occur whenever the driving force exceeds the resisting force.

The multiphase Darcy flow law is the most commonly accepted concept to achieve full physical modelling of fluid flow. The Darcy equation describes the relationship between fluid flow, permeability and pressure gradients (potential gradients) in porous media (e.g., Bear 1972). The Darcy flow concept assumes that flow velocities are controlled by permeability and fluid viscosity functions and it is therefore considered the best way to describe the rates of fluid flow in a physically correct deterministic fashion. The multi-component formulation of these concepts can take into account the dissolution–exsolution behaviour of individual chemical components/compound classes into and out of the phases. This is an important prerequisite because pressure and temperature conditions

· · · ·

full Darcy flow modelling Ortho-contouring Ray-tracing Invasion percolation

Transport processes

Ortho-contouring The simplest method of hydrocarbon flow modelling is ortho-contouring, which focuses on assessing the principal flow directions rather than explicitly modelling the fluid flow process. Contours are constructed at angles of 90 degrees to the strike of a carrier top map. The method is attractive to geologists because it can be performed both manually and by computerised methods. It neglects, however, effects of hydrodynamics (overpressure variations) and carrier entry pressure variations (e.g., carrier facies changes). This method does not provide any information on fluid flow volumes or masses and should therefore not be used for detailed flow analysis. Ray-tracing The ray-tracing method constructs migration flowpaths based on the top depth map of a carrier bed (Sylta 1993). The underlying assumption is that hydrocarbon flow rates and therefore hydrocarbon saturations change very slowly as compared to the duration of simulation time-steps and that migration can therefore be modelled as a steady state process. The thickness of the migration pathways is calculated from the flow rates, carrier bed dips and permeability. Ray-tracing methods are typically applied to migration within highly permeable carrier beds (Skjervøy and Sylta 1993). The method operates on high-resolution grids and requires only short computation times. Figure 6.2.17. Normal fault in a deltaic sand-mudstone sequence, airport road ourcrop, Sarawak, Malaysia. Note the continuous seam of clay in the fault gouge, making this section of the fault a good lateral seal. However, the processes which led to the formation of this clay smear are not clear

usually vary during secondary migration. This method is probably the best when it comes to dealing with the dynamics of petroleum migration, including, in particular, the time required for migration. However, it is also the calculation method which is computationally most demanding. Therefore, several approaches have been investigated to simplify simulation algorithms.

Invasion percolation In the invasion percolation method, flow is modelled based on a pore “entry pressure” (threshold pressure) and petroleum saturation is often assumed to be constant. Migration between adjacent cells is computed by searching for the neighbouring cell with the lowest resistance to flow. All flow is then modelled to occur into this neighbouring cell. Repeated application of the algorithm results in the identification of a “backbone” along which essentially all migration of the non-wetting fluid (petroleum) occurs. This method is extremely efficient due to the concentraFigure 6.2.18. Schematic sketch of the stress evolution, in Mohr space, of a point (A) in the clay layer in a releasing bend of a normal fault

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Figure 6.2.19. Fluid regimes, meteoric, compactional and thermobaric, in sedimentary basins including some HC fluid types (modified after Galloway 1984 and Harrison and Tempel 1993)

tion of computing resources on the calculation of the hydrocarbon migration backbone. The main advantage of the method is the computational speed which makes it possible to create transport simulations on basin-scale models having seismic property resolutions.

thermal or thermohaline convection and seismic pumping may act as driving mechanisms for fluid flow in sedimentary basins (Bjørlykke 1993; Garven 1995; Harrison and Tempel 1993; Sharp et al. 2001; Sibson et al. 2003; Wood and Hewett 1984; Gaupp et al. this volume).

All four approaches outlined above are currently in use in commercial basin modelling packages. Applications have, however, shown that in some cases one of these methods tends to provide better results (in terms of reproducing observations) than the others. Combined (“hybrid”) methods, e.g., with Darcy flow assumed in low permeability rocks and ray-tracing in carrier systems, are also available.

Meteoric water flow is a consequence of the infiltration of rainfall into the subsurface. Its direction and intensity depends on the topography (hydraulic head) and the permeability, lateral extension and connectivity of the conducting rock layers (aquifers). Therefore, meteoric fluid flow is most important at shallow depths or in uplifted parts of a basin. At topographic highs along marine basins, however, meteoric waters can displace connate fluids down to depths of 3 km (Harrison and Temple 1993). For example, freshwater springs discharge at depths of several hundred meters beneath sea level in many coastal regions.

Expulsion of petroleum from source rocks is a slow process. Expulsion rates vary between a few grams to a few kilograms of hydrocarbons per m3 source rock per million years. If the rate of build-up of the driving force (i.e., an overpressure gradient towards the resisting pores) is very low, it can be shown that viscous forces are negligible. Then, the fluid flow is dominated by capillary forces and the result is an extremely focused flow with very narrow pathways in sandstones (<1 m). These arguments have led England et al. (1987) to describe the fluid flow pattern as “migration ganglia” or “rivers of oil (and gas)”, describing the very focused and discrete distribution of hydrocarbons in sedimentary basins.

6.2.4 Geological aspects of fluid transport 6.2.4.1 Fluid flow regimes Different flow regimes occur in sedimentary basins; meteoric, compactional and thermobaric (Fig. 6.2.19). In addition,

Some meteoric water systems extend over large distances and transport considerable amounts of water over geological time. Under varying pressure and temperature conditions the solubility of ionic and gaseous species in the water changes. This may result in a redistribution of these species within the hydrological system. Cramer et al. (1999) investigated the potential role of the aquifer transport of dissolved methane in the formation of the giant Urengoy gas field in Western Siberia. Assuming that methane saturation of the water in the Cretaceous aquifer of the West Siberian Basin occurs some 500 km south of the Urengoy field, they concluded that long-range migration in combination with a Neogene uplift and gas exsolution can account for the volumes of isotopically light methane in the giant gas field (Fig. 6.2.20). If overburden pressure increases faster than compaction, overpressure can form in under-compacted zones. Compactional flow occurs when pore fluid is expelled

Transport processes

Figure 6.2.20. Profile across the W. Siberian Basin and piezometric iso-potential lines of the Cretaceous aquifer demonstrating the concept of long-distance aquifer transport of methane into the giant Urengoy gas field (modified after Cramer et al. 1999)

from the over-pressured zone. As pore fluids move from compacting fine-grained sediments into more permeable coarser-grained sediments, overpressures will dissipate within relatively short geological times and the hydrostatic pressure regime is approached. Consequently, compaction driven flow is generally considered only as a localised process and is of lesser importance in regional fluid movement. The compactional flow regime reaches from the surface down to a depth of approximately 4 km (Harrison and Temple 1993). Thermal gradients increase at greater burial depth and can induce thermal convection through permeable strata. In particular in areas with dominant salt structures, which are excellent heat conductors, the temperature field becomes disturbed. In addition, salinity gradients can occur in the vicinity of salt domes triggering thermohaline convection as described by Sharp et al. (2001) or Magri et al. (this volume), which can drive fluid flow through large portions of the sedimentary section. Fluid convection not only transports heat and energy but also dissolves species which drive diagenetic processes. Repetitive cycling of brine within an essentially closed system may result in considerable redistribution of matter by dissolution and precipitation/exsolution processes. Because such convective flow can develop on the scale of entire basins, it is an important process in diagenesis and emplacement of petroleum. In the deepest portion of sedimentary basins, thermobaric flow can occur. Fluids move in response to pressure changes generated by the release of mineral-bound water, the generation of hydrocarbons, increased heat flow and/ or continuous lithostatic loading (Galloway 1984). The

dehydration of clay minerals like the illitisation of smectites can release significant amounts of water to the pore fluid. Increased heat flow may cause volume expansion of pore water also raising static pressure. However, low volumes and permeabilities at great depth restrict water movement over geological time spans (Galloway 1984) and small fluid volume movements rapidly dissipate the overpressure generated in the thermobaric regime. Sibson (2003) defined the term seismic pumping as fluid flow driven by repeated episodes of dilation and contraction near and within seismically active fault zones. Dilation drives fluids towards newly opened cracks and dilated pore space. Because of the low permeability of basement rocks, groundwater migrates slowly to faults, resulting in continued fluid flow for weeks following an earthquake. Within sedimentary basins different pressure compartments separated by impermeable barriers preventing pressure re-equilibration can exist. Big lateral pressure differences occurring across faults (Mann and Mackenzie 1990) indicate that faults can act as a lateral seal for the pressure compartments. Over- and underlying sealing layers act as vertical boundaries for the pressure compartments (Borge 2002; Wangen 2001a,b). Distinct pressuredepth gradients in the pre-Cretaceous sandstones in the Viking and the Central Graben suggest pressure compartmentalisation of the sandstones (Darby et al. 1996; Moss et al. 2003; di Primio et al. this volume). Osborne and Swarbrick (1997) evaluated the different mechanisms for generating overpressure in sedimentary basins and concluded that disequilibrium compaction is the most feasible overpressuring mechanism in thick and rapidly deposited

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mudstone sequences. This is supported by the observation that overpressured fluids are often found in young sedimentary sequences that are characterised by rapid sedimentation of thick shale-sand successions (Harrison and Summa 1991). No consensus exists on concepts of static and dynamic conditions in overpressured systems (Moss et al. 2003). The static school considers overpressure as a static system of isolated fluid compartments being permanently sealed off from pressure communication with neighbouring rocks (Bradley 1975; Powley 1990; Hunt 1995; Bradley and Powley 1994). Alternatively, the dynamic school (Bethke 1986; Bredehoeft et al. 1994; Neuzil 1994) considers overpressure formation a transient phenomenon which results from fluid flow in a dynamic, hydraulically continuous system (Moss et al. 2003). Therefore, fluid flow is never completely sealed off and overpressure bleeds off on geological time scales of million of years (Deming 1994; Neuzil 1994). Fluid flow and diagenesis Aqueous phase transport of dissolved chemical species through the pore space of sedimentary rocks may cause cement precipitation or leaching of mineral grains. Diffusive transport is usually relevant only over short distances and/or over long time spans. With respect to their chemical composition, different fluid environments and sources can be distinguished: meteoric water, formation water and fluids released from mineral bound water. Meteoric water is in general an acidic fluid with low salinity which infiltrates the rocks from the surface and reacts with the minerals. It replaces or mixes with existing connate water and thus changes the chemical composition of the fluid and initiates early diagenetic alteration of the rocks. The solute load of the fluid increases with geological time. In the North Sea region, meteoric fluid flow is commonly associated with shallow depth and early diagenetic processes (Haszeldine et al. 2000). Dissolution of K-feldspar and the formation of kaolinite is frequently interpreted as an indicator of inflow of meteoric water into sandstones (e. g., Bjørklykke 1994; Brosse et al. 2003). The kaolinitisation of K-feldspar is not an isochemical reaction because potassium and silica form as reaction

products that have to be removed by pore fluid flow if this process is to proceed. At greater depth or when the temperature field is disturbed by geological structures such as diapirs, thermal convection can transport fluids through the rock, moving them through changing temperature fields or bringing them into contact with rocks of different mineral composition. The new conditions change the existing chemistry of the fluid rock system which will cause mineral reaction thus driving the diagenetic evolution of the rock. The capacity for solute transport by compaction driven flow is generally very low and the resulting effect on diagenetic processes is limited. If, however, compactional flow is focused by faults this can allow an efficient transfer of overpressured fluids of distinctively different composition into regions of lower temperature and normal pressure which may lead to significant mineral precipitation near the fracture zone. A study of the diagenetic evolution of Permian Sandstones in North Germany (Gaupp et al. 1993; Platt 1993) indicates that K-feldspar dissolution and illite formation are closely linked to fault zones indicating that faults acted as conduits for fluid flow which triggered subsequent mineral reactions. A study of the diagenetic evolution of Gulf Coast mudrocks (Land and Milliken 2000) shows a loss of silica and carbonate and an import of potassium. No clear answer is given with respect to the relevant transport mechanism but Land and Milliken (2000) consider diffusive transport unlikely because chemical gradients in the large-scale sand-mud system are insufficient to drive diffusion. Alternatively, thermal convection or transport along fractures are considered possible but unproven transport processes. Although there is some controversy as to the exact extent and efficiency of the different driving mechanisms for fluid flow in sedimentary basins (e.g., Haszeldine et al. 2000; Bjørlykke 1993; Land and Milliken 2000), fluid flow is an efficient process for transporting reactants through the pore space of rocks. It is, therefore, an absolutely essential process which drives diagenetic processes in sedimentary basins. Just which transport process, fluid flow or diffusion, controls diagenesis depends on the geological conditions of the system and it may change over time.

6

Chapter 6.3

Fluid-rock interactions R. Schöner · V. Lüders · R. Ondrak · R. Gaupp · P. Möller

6.3.1 Introduction The properties of sediments and particularly of lithified sedimentary rocks depend on the constituent solid materials, their physical and chemical assembly, the type and distribution of pore fluids, and the interaction between solid and fluid components. For the reconstruction of the compositional evolution of fluids, their migration, and the prediction of sediment properties for economic purposes, a thorough understanding of the history of fluidrock interactions is essential. This implies that interaction between solids and fluids must be deciphered from traces of reactions left in the solid matter and findings obtained should be reconciled chronologically with the geological evolution of the basin. It is far beyond the scope of this chapter to give a comprehensive overview of all aspects of fluid-rock interactions from the surface to deep regions. Instead, the emphasis will be on the chemical evolution of deep brines, the reconstruction of their origin and their temporary composition during stages of fluid-rock interactions occurring in clastic petroleum reservoir rocks. The examples will focus especially on the very well studied Late Palaeozoic sediments of the North German Basin (NGB), the central part of the Central European Basin system (CEBS). Fluid evolution above the Zechstein evaporites is described by Magri et al. (2005b, this volume), Möller et al. (2007b), Tesmer et al. (2007) and Gaupp et al. (this volume). For further approaches and references to water-rock interactions we refer to, e.g., Stober and Bucher (2002) and Wanty and Seal (2004).

6.3.2 Evolution of deep brines 6.3.2.1 Origin of saline brines Most sedimentary rocks contain connate pore water from the time of deposition, but the composition of this water commonly has been altered during subsequent burial and basin evolution owing to physical and chemical processes. The controlling mechanisms of fluid evolution are still debated after decades of economic and scientific investigation. Deep basinal brines may contain a load of total dissolved solids much higher than surface marine water.

The dominant anion in most deep brines is chloride, the major cations are sodium, magnesium, calcium and potassium (see also Gaupp et al. this volume). In most cases, Na-(Mg)-Cl or Na-(K)-Cl brines develop into Na-Ca-Cl and ultimately into Ca-Na-Cl or even Ca-Na-K-Cl brines. Carpenter (1978) suggested that the chemical composition of many deep basinal brines resembles evaporated seawater, which interacted with the sedimentary rocks and/or has subsequently mixed with seawater or meteoric water. Clayton et al. (1966) published isotope data which strongly indicate that brines in several basins originated from continental, meteoric water and not from seawater (see also Gaupp et al. this volume). Any type of water may evolve to saline brines by subsurface evaporite dissolution. Nevertheless, the composition of most deep brines cannot be explained by evaporation of water and salt dissolution alone, but suggests modification by reactions with minerals in the sedimentary rocks. Kharaka and Thordsen (1992), Hanor (1994a,b) and Rosenthal (1997) summarised the results of geochemical, isotopic and hydrodynamic studies on the origin and compositional evolution of brines in sedimentary basins. They emphasised that a variety of physical and geochemical processes can modify the composition of saline subsurface brines. Evaporation of seawater evidently leads to Mg-Cl brines with salinities up to 400 g/l (e.g., White 1965; Carpenter 1978; Warren this volume). The removal of water and successive precipitation of carbonate, sulfate and salt minerals produces predictable changes in ionic ratios during evaporation. Relatively low volumes of CaCO3 precipitate from water concentrated to about twice the salinity of sea water. Up to salinities of about 100 g/l all dissolved species evolve at similar rates. With continued evaporation Ca2+ is removed from solution by precipitation of gypsum. Concentrations above approximately 340 g/l result in crystallisation of halite and depletion of Na+, whereas Mg2+, K+, SO42-, and Br- increase significantly. The geochemical signature of the residual fluid evolves from Na-Cl to Mg-Cl. Typical examples are brines developing within isolated basins repeatedly open to oceans (for example, during the Zechstein in the CEBS). Evaporation of continental waters produces brines with less predictable composition as the initial water chemis-

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try varies for individual basins. The majority of evaporated continental pore waters are dominated by Na+, Ca2+ and/or Mg2+ and may contain Cl-, SO42- or HCO3--CO32- as the dominant anion (Eugster and Jones 1979; Jones and Deocampo 2003). Typical brines which developed from fresh water are those of modern terrestrial terminal lakes such as the Dead Sea (Middle East), Aral Sea (Central Asia), Great Salt Lake (USA) and Lake Eyre (Australia). Similar conditions must be assumed for the large Permian Rotliegend terminal lake in the CEBS, except for the uppermost Rotliegend, which already shows an input of seawater (see also Stollhofen et al. this volume). Subsurface dissolution of evaporites (halite, anhydrite, gypsum, magnesium and potassium salts) is very effective in producing high-salinity brines (Bennett and Hanor 1987; Bein and Dutton 1993). Likewise, leaching of brines (enclosed in pore space or trapped as fluid inclusions in minerals) may increase the amount of the dissolved salt load. Saline brines derived from halite dissolution have generally high Cl/Br ratios, because little bromine is incorporated into the halite lattice due to its large ionic radius. The SO4/Cl ratio varies depending on the amount of anhydrite and halite dissolution. Diffusion driven by high concentration gradients, and convection driven by density differences, can cause transport of dissolved solids over very large distances around salt domes. Additional processes have been suggested to account for high salinities in subsurface brines. Freezing of seawater removes H2O as ice and thus concentrates dissolved solids in the residual water (e.g., Herut et al. 1990). This process may be important in polar regions and during glacial periods. Filtration of charged ions at semi-permeable shales during pressure-driven cross-formational fluid flow has been discussed as another mechanism to increase salinity (e.g., Kharaka and Berry 1973; Graf 1982). However, the significance of shale-membrane filtration for the evolution of brines in sedimentary basins is discussed controversially, since there is no convincing field evidence (Hanor 1994a).

6.3.2.2 Fluid-rock interactions modifying the composition of brines Most diagenetic reactions are not isochemical mineral transformations, but release specific ions and/or remove other ions from the formation water. Many water-rock interactions mainly change the composition of the formation water rather than the salinity. Exceptions are dissolution or precipitation of chlorides and sulfates as well as reactions releasing H2O, for example dehydration of clay minerals. A number of reactions can explain why a saline water primary enriched in Mg, evolves into a Ca-Cl brine. Reaction of pore fluids with limestones can account

for a significant decrease in the Mg/Ca ratio. Fluid flow through chalk beds leads to Ca-Cl water composition by dissolution of CaCO3 and ion exchange with smectite, by which Ca2+ is released and Na+ is consumed (Rosenthal 1997). Alternatively, diagenetic dolomitisation of calcium carbonates releases Ca2+ and removes Mg2+ from the formation water (e.g., Goldsmith and Graf 1958; White 1965): 2 CaCO3 + Mg2+ → CaMg(CO3)2 + Ca2+ calcite dolomite

(6.3.1)

Albitisation of the anorthite component of plagioclase occuring in clastic sediments or igneous rocks yields Ca2+ at the expense of Na+. This process takes place during burial diagenesis and can account, for example, for the composition of Ca-Cl brines in Cenozoic clastic sediments of the Gulf of Mexico (Fisher and Boles 1990; Perez and Boles 2005). Several albitisation reactions have been proposed, e.g., by Land (1984, Eq. 6.3.2) and by Boles (1982, Eq. 6.3.3): CaAl2Si2O8 + 2 Na+ + 4 H4SiO4 → 2 NaAlSi3O8 + Ca2+ + anorthite albite 8 H2O (6.3.2) CaAl2Si2O8 + 2 SiO2 + Na+ + H+ → NaAlSi3O8 + 0.5 + anorthite quartz albite Al2Si2O5(OH)4 + Ca2+ (6.3.3) kaolinite Albitisation of potassium feldspar is an important source of K+ during deep burial diagenesis (Eq. 6.3.4). Field studies as well as thermodynamic and kinetic modelling suggest that albitisation of K-feldspar is enhanced by increasing temperature, and is probably linked to clay mineral reactions that consume K+ (Aagaard et al. 1990; Ben Baccar et al. 1993; Krooss et al. this volume). KAlSi3O8 + Na+ → NaAlSi3O8 + K+ orthoclase albite

(6.3.4)

Chloritisation of biotite and Ca-rich feldspars in igneous rocks and their clastic debris in sediments (e.g., in the Permian Rotliegend volcanic and intrusive rocks and associated arkoses/litharenites in the CEBS) takes up Mg2+ and releases Ca2+ (Eqs. 6.3.5, 6.3.6). Hence, chloritisation may contribute to the formation of Ca-Cl brines. Additionally, equation 6.3.5 suggests that chloritisation could also decrease the Na+/K+ ratio in the fluid. 2 K(Mg)3[AlSi3O10(OH)2] + Mg2Si2O6 + 2 Mg2+ + biotite pyroxene 2 CaAl2Si2O8 + 6 H2O + 2 Na+ → 2 (Mg2Al)[AlSi3O10 anorthite clinochlore (OH)2]·Mg3(OH)6 + 2 NaAlSi3O8 + 2 Ca2+ + 2 K+ albite (6.3.5)

Fluid-rock interactions

CaAl2Si2O8 + 2 Mg2Si2O6 + 10 H2O + Mg2+ → anorthite pyroxene (Mg2Al)[AlSi3O10(OH)2]·Mg3(OH)6 + Ca2+ + 3 H4SiO4 clinochlore (6.3.6) Many different clay mineral reactions may take place from early diagenesis through deep burial and have considerable potential to modify the composition of formation waters. One important reaction is the illitisation of smectites in sediments, which consumes K+ and releases variable amounts of Na+, Ca2+, Mg2+, Fe2+, Si4+ and water, depending on smectite composition (e.g., Boles and Franks 1979). The reaction is a progressive process that takes place during burial, and is mainly controlled by temperature, time and concentration of potassium in solution (e.g., Meunier and Velde 2004). There are several possible summary reactions for the conversion of smectite into illite. Worden and Morad (2003) provided an equation conserving Al, which is considered as largely immobile in most diagenetic environments: 1.58 K0.1Na0.1Ca0.2Mg0.4Fe0.4Al1.4Si3.8O10(OH)2·H2O + dioctahedral smectite 0.393 K+ + 1.242 H+ → K0.55Mg0.2Fe0.15Al2.2Si3.5O10(OH)2 illite + 0.16 Na+ + 0.31 Ca2+ + 0.43 Mg2+ + 0.24 Fe2O3 + 2.47 SiO2 + 2.86 H2O (6.3.7) The consumption of potassium may be accommodated by the breakdown of detrital K-feldspar. The components released by the smectite-illite conversion are commonly consumed during further diagenetic reactions. Other important clay mineral reactions are the chloritisation of smectite and/or early diagenetic Mg-Fe-rich clay minerals, the growth of kaolinite or illite at the expense of feldspar and the illitisation or chloritisation of kaolinite (for review, see Worden and Morad 2003). These reactions are by no means exclusive, but resemble processes occurring in many sedimentary basins during burial. The examples given above do not account for mixing or replacement of fluids in active fluid flow regimes. In nature, however, mixing of fluids from different stratigraphic units and lithologies, or replacement of formation waters, is very common and causes a range of mineral-fluid interactions. Examples are episodic influx of ascending water generated during compaction of shales (e.g., Cartwright 1994), infiltration of meteoric water (e.g., Bjørlykke et al. 1989), mixing of water from continental clastic and evaporite lithologies (e.g., Purvis 1992), migration of organic maturation products from source to reservoir rocks (e.g., Surdam et al. 1989) and influx of hydrothermal fluids from basement rocks (e.g., Lüders et al. 2005). Some of these more complex fluid-rock interactions are addressed below in example of the CEBS.

6.3.3 Palaeo-fluid reconstruction 6.3.3.1 Methods of palaeo-fluid reconstruction The reconstruction of fluid evolution from the surface to deep burial must take into account compositional evolution of fluids with increasing temperature and pressure, fluid-rock interactions and stages of fluid migration. The degree of fluid-rock interaction related to stages of fluid migration is closely linked with fluid flow through networks of microfractures or focussed flow along permeable fault zones. Type, succession, age and chemistry of authigenic minerals and their fluid inclusions, as well as the chemistry of formation waters, reflect the evolution of basinal brines. However, the preserved evidence of palaeo-fluid composition over geological time is neither complete nor continuous, but is highly selective, and authigenic minerals that can be dated are limited. Palaeo-fluid reconstructions always start with (i) the petrographic analysis of the paragenetic sequence of diagenetic processes and products (that already involves interpretations). Diagenetic studies should include the analysis of minor and trace element chemistry of authigenic minerals by (ii) electron microprobe and/or (iii) cathodoluminescence microscopy plus spectroscopy (e.g., Richter et al. 2003). These chemical data have proven useful to determine the nature of dissolved species and the redox potential at the time of mineral precipitation. Based on these investigations, several methods assist the reconstruction of palaeo-fluids: 1. Fluid inclusion microthermometry 2. Isotope chemistry of authigenic minerals 3. Geochronological dating of K-bearing authigenic minerals (illite, K-feldspar) 4. Thermochronology (apatite fission track ages, U/ThHe dating) 5. Numerical modelling of fluid-rock interactions 6. Integrative evaluation of data and modelling results (including geological/structural history, thermal and organic maturation evolution, fluid flow and reactive flow modelling) Fluid inclusions Primary fluid inclusions hosted in diagenetic cements and vein minerals trace flow pathways of palaeo-fluids and can, therefore, record important information for the reconstruction of the thermal history and fluid evolution within a sedimentary basin. As pointed out by Goldstein and Reynolds (1994), an important prerequisite for a successful fluid inclusions study in sedimentary terrains is

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R. Schöner · V. Lüders · R. Ondrak · R. Gaupp · P. Möller Figure 6.3.1. Sketches of primary and secondary fluid inclusions (FI, further explanation see box 6.3.1) in detrital quartz and authigenic quartz overgrowths

the careful analysis of paragenetic relationships of diagenetic minerals that formed at different stages of subsidence or uplift together with the petrography of the fluid inclusion assemblages therein (Fig. 6.3.1). Although fluid inclusions may undergo post entrapment alteration due to changes in temperature and/or pressure (Roedder 1984; Goldstein and Reynolds 1994), under the best circumstances the sources and composition of fluids associated with hydrocarbon migration can be traced, and pressuretemperature-composition (P-T-X) conditions for fluid inclusion entrapment can be reconstructed (e.g., Goldstein and Reynolds 1994; Dubessy et al. 2001; Thiéry et al. 2002). The chemical composition of fluid inclusions can be determined by several means (Table 6.3.1), among which microthermometry and laser Raman spectroscopy are the most frequently applied micro-analytical methods for the determination of salinity, gas composition and homoge-

nisation temperature (minimum temperature of mineral formation) of fluid inclusions in diagenetic minerals (Box 6.3.1 Fig. 2). In cases where gas-bearing inclusions are trapped along with aqueous fluid inclusions the P-T conditions of fluid entrapment can be estimated via isochore construction as illustrated in figure 6.3.2. Analysis of anion and cation content in fluid inclusions provides direct evidence on the origin of salinity in fluid inclusions as well as the possible degree of water-rock interaction. Figure 6.3.3 shows crush-leach bulk analyses of fluid inclusions hosted in various minerals (quartz, carbonates, anhydrite and fluorite) from fissure mineralisation of Palaeozoic units from the NGB. The Cl/Br ratios are highly variable and suggest different origins of salinity, i.e., either (sea)water evaporation or dissolution of halite. Irrespective of this, the fluid inclusions also show variable Na/K molar ratios. In general, fluids showing low Na/K ratios (< 30) have undergone extensive water-rock

Table 6.3.1. Methods to determine the chemical composition of fluid inclusions Microthermometry

e.g., Roedder (1984), Goldstein and Reynolds (1994)

Crush-leach bulk analysis

Banks and Yardley (1992)

Laser ablation ICP mass spectrometry

Günther et al. (1998), Heinrich et al. (2003), Gagnon et al. (2003)

Proton induced X-ray emission (PIXE) analysis Synchrotron radiation induced X-ray fluorescence microanalysis

Anderson et al. (1989), Kamenetsky et al. (2002), Anderson and Mayanovic (2003) e.g., Mavrogenes et al. (1995), Ménez et al. (2002)

Fourier transform infrared (FTIR) spectroscopy

e.g., Pironon et al. (2001)

Laser Raman spectroscopy

Burruss (2003) and references therein

Fluid-rock interactions Figure 6.3.2. Idealised P-T phase diagram of co-genetic aqueous and gaseous fluid inclusions. Isochores for fluid inclusions can be calculated with special computer programs (e.g., Brown and Hagemann 1994; Bakker 2003; Bakker and Brown 2003). The intersection of the isochors of co-genetic aqueous and gaseous inclusions in a mineral defines the minimum P-T conditions of fluid entrapment

interactions as indicated by significant loss of Na (see above). Isotopic signatures of authigenic minerals The determination of light stable isotopic compositions (H, C, O, and S) and strontium isotopic composition (87Sr/ 86 Sr) in diagenetic minerals can provide information on the source of the fluids that formed them, formation temperatures and physico-chemical conditions during mineral precipitation (e.g., O’Neil et al. 1969; Robinson 1975; Veizer and Hoefs 1976; Ohmoto and Rye 1979; Taylor 1987, 1997; Ziegler 2006).

The carbon isotopic composition of carbonates and bicarbonate in nature can vary extremely between -64.5‰ and 37.5‰ (Deuser 1970; Claypool et al. 1985; Hoefs 1997) and depends on temperature, pH, fO2 and/or the ionic strength (I) as well as the relative proportions of carbon species present in the fluid (Fig. 6.3.4). Formation waters may contain H2CO3 (= aqueous CO2), HCO3-, CO32- or CH4. However, since amounts of CO32- or CH4 are generally very low in carbonate-forming fluids (Ohmoto and Rye 1979), the carbon isotopic fractionation between carbonate and fluid is predominantly restricted to the proportions of H2CO3 (including CO2 which is asFigure 6.3.3. Cl/Br vs. Na/K diagram for fluid inclusions hosted in fissure minerals. For details see text

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Box 6.3.1 Fluid Inclusions Fluid inclusions are micron-scale liquid and/or gas filled cavities in minerals that either formed on imperfections on the crystal surface during crystallisation (primary inclusions, e.g., Roedder 1984) or on sealed microcracks due to post crystallisation deformation processes (secondary inclusions). Fluid inclusions occur in nearly every environment and provide direct clues to the composition of mineral-forming fluids and temperature of crystallisation. They are highly variable in shape, size and the composition of the trapped fluids. In low-temperature environments fluid inclusions trapped in cements or vein minerals typically contain only a liquid phase of the mineral-forming fluid. Minerals that formed from hydrothermal or magmatic fluids often contain two or even more phases. Detailed studies of fluid inclusions using various techniques (Table 6.3.1) provide information on pressure-temperature-composition (P-T-X) conditions existing during mineral formation or post crystallisation hydrothermal overprinting.

Figure 1. Fluid inclusions in fissure minerals from boreholes in the CEBS, northern Germany. Left: cluster of co-genetically trapped aqueous two-phase inclusions and dark CH4-CO2-bearing gas inclusions in quartz hosted by Carboniferous sandstones. Right: high-salinity aqueous two-phase fluid inclusion with a halite daughter crystal in fluorite from a fissure in Zechstein carbonate (Ca2) Figure 2. Idealised P-T phase diagram for a two-phase H2ONaCl fluid inclusion. The fluid inclusion was trapped as a monophase (liquid) inclusion at maximal burial P-T conditions (grey ellipse). During cooling and uplift the inclusion follows the slope of its specific isochore until it reaches the liquidvapour phase boundary where it nucleates a bubble. At room temperature the inclusion shows a liquid phase and a vapour bubble. During microthermometric heating runs the inclusion is heated until the vapour bubbles disappears. The measured homogenisation temperature Th is the minimum trapping temperature of the inclusion. Pressure estimates derived from geological information (e.g., burial curves) may be used to estimate the true temperatures of fluid entrapment (Ttrap)

sumed to behave like H2CO3 isotopically) and/or HCO3-: Carbonate

Carbonate

HCO -

Δ13CFluid = Δ13CCO2 - XHCO3- • Δ13CCO2 3

(6.3.8)

Δ13Cba is the carbon isotopic fractionation between component a and b. XHCO3- is the mole fraction of HCO3- in the fluid with respect to apparent H2CO3 (including dissolved CO2 and H2CO3) and is linked by the following reaction: H2CO3 aq. ↔ HCO3- + H+ (6.3.9)

The oxygen isotopic composition of carbonates is mainly controlled by the isotopic composition of the formation water even though the proportion of dissolved CO2 cannot be completely neglected. The oxygen isotopic composition of carbonates therefore can be described by: Carbonate

CO

δ18O Carbonate = δ18O Fluid + Δ18OH2O - XCO2 • Δ18OH2O2 (6.3.10) XCO2 is the mole fraction of CO2 in the fluid relative to H2O.

Fluid-rock interactions

In deep basin fluids changes in temperature may be a common cause of observed carbon and oxygen fractionation in carbonates. At temperatures below 200 °C carbon isotopic fractionation between calcite and HCO3- becomes nearly constant (Ohmoto and Rye 1979). Therefore, variations in carbon isotopic compositions of sedimentary and diagenetic calcites reflect the carbon isotopic composition of the fluid, i.e., under the best circumstances also the fluid source (Fig. 6.3.4). The oxygen isotopic composition of calcite, however, increases with decreasing temperature in all relevant mineral-fluid systems (O’Neil et al. 1969). This allows the calculation of the oxygen isotopic composition of the fluid, provided that the temperature of mineral growth is constrained by other methods, for example by fluid inclusion microthermometry. In a similar way, oxygen isotopic compositions are used to study authigenic sulfate, quartz and clay minerals. An example of how carbon and oxygen isotopic compositions can be used to determine the nature and evolution of fluids in sedimentary basins is shown in figure 6.3.5. The δ18O and δ13C values of fissure calcites in Carboniferous and Rotliegend clastic rocks of the NGB are typical of deep basinal (hydrothermal) calcites. A significant organic carbon compound in the calcite-forming fluids can be excluded considering the range of δ13C values (Fig. 6.3.4). The majority of calcites hosted by Permian volcanic rocks show similar δ18O and δ13C values to fissure calcites in Permian or Carboniferous clastic rocks. Some calcites yield heavier oxygen, but constant carbon isotopic composition, which can be explained by rock-buffered interactions of the fluids and the volcanic host rocks

(trend B in figure 6.3.5). A group of calcites from miarolic cavities shows light δ13C and high δ18O typical for calcites that have been remobilised (trend A in figure 6.3.5). Very similar carbon and oxygen isotopic compositions, as well as similar salinity and homogenisation temperatures of fluid inclusions, have been observed for calcites from hydrothermal vein mineralisations in the Harz Mountains that deposited from highly saline brines at temperatures of about 200 °C (Fig. 6.3.5; Lüders and Möller 1992; Lüders et al. 2005). Hence, it seems likely that the fissure calcites in Rotliegend and Carboniferous rocks precipitated from ascending deep basinal (hydrothermal) fluids that interacted with the host rocks to varying degrees. A different situation is recorded in diagenetic calcites hosted by Zechstein carbonates. Calcites showing a shift in δ18O between +18 and +25‰ and δ13C between -5 and +5‰ may be interpreted to be the product of temperature-controlled isotope re-equilibration (Bottinga 1969; O’Neil et al. 1969; Fischer et al. 2006). This suggests that they precipitated from heated Zechstein pore fluids that achieved equilibrium with the surrounding carbonates. Some late diagenetic calcites showing δ13C below -10‰ may contain minor amounts of organic carbon that mixed with marine waters (Fischer et al. 2006). Sulfur isotopic compositions of authigenic anhydrite and barite may yield highly variable δ34S values. For example, Lüders et al. (2005) reported δ34S values ranging from about -20‰ to almost +24‰ for sulfates from fissure mineralisation in Carboniferous and Permian rocks. Furthermore, fluid inclusions in anhydrite show very heterogeneous ranges for Th and Tm ice (Lüders et al. 2005) Figure 6.3.4. Carbon isotopic composition of carbon compounds (after Hoefs 1997) and calcites from fissures from various occurrences in the North German Basin

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Figure 6.3.5. δ18O vs. δ13C relationship in fissure calcites hosted by different Palaeozoic stratigraphic units of the NGB. The ranges of δ18O and δ13C values of hydrothermal calcites from the St. Andreasberg and Upper Harz Mts. mining districts (Lüders and Möller 1992) are shown as blue ellipses for comparison

and suggest that anhydrite formation in the North German Basin cannot be related to a uniform fluid source. Instead, the great range of sulfur isotopic composition in the studied samples of anhydrite either indicate different sources for sulfate (i.e., oxidation and re-deposition of dissolved biogenic sulfide, formation waters, and/or mixing of SO42--rich formation waters with variable amounts of dissolved biogenic sulfide) or that anhydrite precipitated under quite different physico-chemical conditions. Sulfur isotopic fractionation in formation waters is a function of temperature, pH, ƒO2, isotopic composition of the fluid, and/or dissolved element species in the fluid (Ohmoto 1972; Ohmoto and Rye 1979). The δ34S values of seawater varied significantly over earth history, with maximum values of +35‰ in the early Cambrian and minimum values below +10‰ in the Permian (Holser 1977). Microbial processes, especially dissimilatory sulfate reduction, produce sulfides strongly depleted in δ34S

relative to seawater. Changes in ƒO2 and/or pH can cause large variations in sulfur isotopic composition of sulfides and sulfates assuming a uniform fluid source (Ohmoto and Lasaga 1982). Nevertheless, δ34S values of authigenic anhydrite and barite mineralisation may help to identify different sources of sulfate. Strontium isotopic ratios (87Sr/86Sr) of minerals directly reflect the ratios of the fluid from which they precipitated, because the very similar mass and size of the two isotopes does not cause significant fractionation effects (Faure 1986). 87Sr/86Sr in seawater, which is considered as homogeneous reservoir, fluctuated from about 0.707 to 0.709 over Phanerozoic times (Burke et al. 1982). The lowest ratios were reached in Permian and Jurassic times. In continental basins the strontium isotopic composition of surface water reflects the geology of the drainage area. Waters derived from igneous and metamorphic basement

Fluid-rock interactions

Figure 6.3.6. Abundance of sulfur isotopic compositions (left) and sulfur vs. strontium isotopic compositions (right) of anhydrite cements from Rotliegend sandstone reservoirs in northern Germany (Platt 1994). Boxes in both diagrams show typical compositions of early sulfates formed from Zechstein (Claypool et al. 1980) and Rotliegend waters (Holser 1979). Some samples from horst blocks indicate an influence of Zechstein fluids by sulfur and strontium isotopic compositions intermediate between those of Zechstein and Rotliegend waters

rocks, for example, typically have high 87Sr/86Sr ratios in the range of 0.720 (Faure 1986). However, the strontium isotopic composition of subsurface brines may be modified by water-rock interactions and by mixing of compositionally different fluids.

its correlation with geological factors. Under best circumstances constraints can be obtained from apatite fission track analysis (AFTA), (U-Th)/He dating on detrital apatite, and K/Ar or Ar/Ar dating of authigenic illite or K-feldspar.

Combined sulfur and strontium isotopic analyses were used, for example, by Platt (1994) to trace the nature of fluids from which anhydrite cements precipitated in Rotliegend sandstones of the NGB (Fig. 6.3.6). Most analysis show sulfur isotopic compositions similar to early Rotliegend sulfates reported by Holser (1979) which most likely precipitated from continental waters. Some burial diagenetic anhydrite cements from horst blocks show higher δ34S values and lower strontium isotopic ratios than typical early Rotliegend sulfates. This suggests mixing of Rotliegend and Zechstein pore fluids across fault zones confining the horst blocks, where Rotliegend and Zechstein rocks are juxtaposed.

ATFA enables the determination of the time-temperature path if a horizon was cooled after maximum temperature exposure (Fig. 6.3.7; Green et al. 1989; Duddy et al. 1998). This method is used for sedimentary basin thermal history reconstructions. Apatite (U-Th)/He thermochronology is based on progressive resetting of (U-Th)/He ages by heating, with total loss of helium occurring at temperatures around 80 to 90 °C over geological timescales. Using this technique, thermal history reconstructions derived from AFTA and vitrinite reflectance (Littke et al. this volume) can be further refined, allowing improved precision of the dating of palaeo-thermal events at relatively low temperatures (< 80 °C) (e.g., Crowhurst et al. 2002).

Thermochronology and age dating of authigenic minerals

Platy or fibrous illite is one of the most important authigenic minerals that grow within sandstones during burial, and the only one which is commonly used for dating using the K/Ar technique (e.g., Lee et al. 1989; Leveille et al. 1997; Liewig and Clauer 2000). Although the processes controlling illite growth are not yet fully understood, temperature, sandstone composition and fluid chemistry

Absolute ages of diagenetic processes like mineral precipitation, or cooling of a reservoir below defined temperature thresholds, allow the subdivision of the (relative) paragenetic sequence of authigenic minerals and

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R. Schöner · V. Lüders · R. Ondrak · R. Gaupp · P. Möller Figure 6.3.7. Thermal history of a sedimentary unit schematically showing the integration of palaeothermometers. Illite data constrain the burial or heating phase of a basin‘s thermal history, vitrinite reflectance records maximum temperature (Tmax), and apatite fission track analysis (AFTA) constrains timing of uplift and cooling (redrawn from Pevear 1999)

(alkaline, K+ bearing) are generally accepted as major influencing factors (e.g., Clauer and Chaudhury 1992; Ziegler 2006). Case studies have shown that stages of hydrothermal activity in sedimentary basins or structural perturbations (fault gauge formation) were accompanied by illite growth (e.g., Zwingmann et al. 1999; Ylagan et al. 2002). Cessation of illite growth is frequently attributed to geological events such as hydrocarbon filling of sandstone reservoirs (gas or oil charging), which can allow calibration to basin models. In oil or gas filled structures diagenetic processes are likely to be retarded or stopped due to the removal of formation waters, whereas in volumes below palaeo-gas/oil-water contacts diagenetic processes proceed as long as physico-chemical conditions allow. Illite K/Ar ages can thus contribute to the temporal and spatial reconstruction of hydrocarbon charging. Illite ages generally become younger down-section in hydrocarbon wells, but multi-phase illite growth can generate complex spatial patterns of illite morphotypes and K/Ar ages (Lee et al. 1989; Gaupp et al. 1993; Zwingmann et al. 1998). Careful separation of detrital and authigenic K-bearing minerals is a prerequisite for obtaining reliable illite ages (e.g., Pevear 1999; Meunier and Velde 2004). 40Ar/39Ar step-heating of the fine-grained illites (e.g., Dong et al. 2000) provides approximate age indications even from mixtures of detrital and authigenic illitic clays. The use of further isotope systems such as Rb/Sr or Sm/Nd can contribute additionally to the understanding of geochronological ages of authigenic clays in sandstones (Liewig et al. 1987; Zwingmann et al. 1999). 40

Ar/39Ar-dating of authigenic K-feldspar is still rarely used in basin analytical studies or thermochronology (Parsons et al. 1999; Mark et al. 2005), but can contribute to unravelling complex fluid flow and temperature-time-

composition histories of sedimentary basins. This method is particularly promising when combined with fluid inclusion microthermometry and other thermochronological analyses of authigenic minerals. Summary In summary, all analytical tools discussed above have to be considered in order to reconstruct the palaeo-fluid history of a sedimentary basin. In large basins the chemical composition of saline water is successively changed and adjusted to the solutes present in the formations of their residence. Thus, the solutes characterise the surrounding geological formation but the water itself may be of different origin. In most cases, the so-called formation water is not connate water sensu stricto. It may either represent evolved connate water or replacement water of younger age than the formation where it is actually stored. Authigenic minerals and their fluid inclusions preserve at least parts of the temporal and spatial fluid evolution. Dating of authigenic minerals can help to transfer the relative succession of diagenetic events to the geological history of the sedimentary basin. It is important to note that authigenic minerals in the pore space of rocks and authigenic fissure minerals in the same sedimentary unit commonly record different information about temperature and chemistry of fluids that formed them. This is not surprising, because fluids can migrate rapidly along faults or fractures across thick sedimentary sequences without major exchange with the pore fluids. A succession of sandstone cements, for example, represents the evolution of the pore waters, whereas fissure minerals can provide direct evidence of the composition of fluids that migrated episodically through fault systems and that may or may not have affected porous sandstones horizons.

Fluid-rock interactions

6.3.3.2 Synthesis of fluid evolution in the Permian Rotliegend of the North German Basin As an example, we summarise here the results of palaeofluid reconstructions and the evolution of formation waters in Permian Rotliegend reservoirs of the North German Basin. The Rotliegend arkosic red beds, stretching from offshore UK via Netherlands, Germany to southern Poland, are by far the best investigated stratigraphic interval in Central Europe due to their economic importance as major gas reservoirs (e.g., Hetzer and Katzung 1985; Glennie 1990, 2001). Numerous studies of the depositional environment, burial histories, thermal and fluid evolution and diagenesis in relation to reservoir quality evolution have been performed during the last three decades (Glennie 1972; Pye and Krinsley 1986; Gaupp et al. 1993; Leveille et al. 1997; and many others). A wealth of data and conceptual interpretations exist in the files of the hydrocarbon exploration companies, but these are mainly confidential. A brief summary is given here concerning aspects related to fluid flow, fluid chemical evolution and diagenesis. The early, near-surface fluids during Rotliegend deposition in an arid endorheic continental sedimentary basin were dominated by meteoric waters that originated from precipitation at the southern basin margin and possibly from minor rainfall within the basin. Fluid composition was modified by evaporation and precipitation/dissolution of minerals. Near-surface meteoric waters in Rotliegend times were generally oxidising, and caused the break-down of Fe-Mg mineral components such as biotite, pyoxene and magnetite, and the precipitation of iron oxides. Repeated desiccation cycles in semi-arid en-

vironments resulted in early diagenetic ironoxide coatings around detrital grains of sandstones and thus in the formation of red beds (Walker 1967). Fluid composition during early diagenesis was largely controlled by depositional environment. In terminal basins with high evaporation such as the Rotliegend basin in the CEBS, fluids became increasingly more saline, more alkaline and enriched in heavier oxygen isotopes from basin margin towards basin centre. This resulted in the formation of early diagenetic carbonate cements along the basin margins, sulfates and finally halite in more central areas of the basin (Drong 1979; Platt 1994; Fig. 6.3.10). Such early cements may account for up to 40-45 vol.-% of a sandstone and can locally occupy the entire open pore space. Sedimentation of low-permeability rocks such as Rotliegend shales and deposition of Zechstein evaporites, as well as decreasing permeability due to compaction and cementation, hampered and finally inhibited the communication of pore fluids in Rotliegend sandstones with meteoric waters. Pore waters became progressively more saline and enriched in 18O with increasing burial depth and temperature (Fig. 6.3.9; Platt 1994). Compaction of clay-rich lithologies deposited in central parts of the Rotliegend basin resulted in the expulsion of pore waters that migrated upwards via fractures and/or permeable layers. A substantial influence of these advective compactional fluids is assumed especially in sandstone layers interfingering with lacustrine shales (Gaupp et al. 1993). Mixing of compositionally different brines from adjacent lithologies is evident from combined isotopic, fluid inclusion, and petrographic-geochemical data. Lithologically very different formations (Carboniferous Coal Figure 6.3.8. Schematic overview of possible fluid exchange between the geochemically differing Late Palaeozoic formations in the gas play of the CEBS in northern Germany. In cases of hydraulic contacts, mineralogical and petrophysical changes in the affected formations resulted. HC = hydrocarbon, TSR = thermochemical sulfate reduction

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Measures as source rocks, Rotliegend red beds as reservoir and Zechstein evaporites as seal) were stratigraphically and tectonically juxtaposed during Mesozoic structural change (Fig. 6.3.8; Gaupp et al. 1993). Export and import relations of the largely different connate and laterevolved pore waters between the formations led to alteration of mineral composition and to porosity-permeability changes (Gaupp et al. 2005). Platt (1994) suggested that up to 40% of Zechstein fluids mixed with Rotliegend brines close to fault contacts (Fig. 6.3.6), but most anhydrite cements in the Rotliegend reservoirs of northern Germany show no substantial influence of Zechstein waters. The latter observation can be explained by the presence of thick mudflat and playa lake shales, which separate Rotliegend reservoirs from the Zechstein. This is also supported by the fact that the present-day composition of Rotliegend brines differs significantly from that of Zechstein brines (Gaupp et al. this volume). Furthermore, fluid inclusions hosted in Rotliegend cements and fissure minerals differ in their chemical composition from those hosted in Zechstein evaporates and fissures therein (Lüders et al. 2005). Where the Rotliegend is mainly composed of permeable sandstones, for example in the southern North Sea area, infiltration of Zechstein waters probably occurred during the Zechstein transgression (Sullivan et al. 1990; Purvis 1992). Decomposition of organic matter during burial of Carboniferous source rocks resulted in the generation of various maturation products including organic acids, CO2, CH4, and liquid hydrocarbons. Gas generation started in late Carboniferous through Permian times, but the main periods of organic maturation in the area of the large German gas

fields were the Triassic to Jurassic and latest Cretaceous to Cenozoic (Neunzert et al. 1996; Schwarzer and Littke 2007). Early charging of Rotliegend sandstones by methane may have taken place as soon as effective seals had formed. Many present Rotliegend gas reservoirs were filled mainly during the Cenozoic, hence large volumes of Late Palaeozoic to Mesozoic gas must have escaped through the thick Zechstein seal (e.g., Krooss et al. 1995; Littke et al. 1995), probably during periods of tectonic activity. The expulsion of acidic, reducing pore fluids from the Carboniferous into the Rotliegend was responsible for complex mineral reactions in the red bed reservoirs, including leaching of potassium feldspar and growth of authigenic illite as well as locally kaolinite/dickite (see below). The positive oxygen isotopic compositions of these clay minerals and associated late quartz cements suggest that they precipitated from 18O-enriched deep basinal brines (Platt 1993). Figure 6.3.9 summarises the possible path of δ18OSMOW of the formation waters in Rotliegend sandstone reservoirs from North German gas fields as revealed from different isotope analytical investigations. Studies of fluid inclusions in the North German Basin have revealed compelling evidence for multiple events of fluid and gas migration from early diagenesis through stages of Mesozoic burial, Cretaceous inversion and locally subsequent burial in the Tertiary (e.g., Rieken 1988; Reutel et al. 1995; Reutel and Lüders 1998; SchmidtMumm and Wolfgramm 2002, 2004; Lüders et al. 2005). Generally, the fluid systems evolved from a low to medium saline H2O-NaCl±KCl type during the early stage of basin subsidence to a high-salinity H2O-NaCl-CaCl2 or H2O-CaCl2-NaCl type during further burial. The migration Figure 6.3.9. Oxygen isotopic composition of fluids that formed different authigenic minerals in Rotliegend sandstones in gas fields of the North German Basin, and the inferred temperature intervals of mineral formation vs. timing. Temperature indications are derived from fluid inclusion microthermometry and paragenetic sequences, combined with thermal modelling. Q1, Q2a, Q2b are successive generations of authigenic quartz; Cc, Do, An indicate early diagenetic calcite, dolomite and anhydrite, respectively. This composite graph shows the positive δ18O composition of highly evolved brines that precipitated illite and associated quartz Q2a and Q2b (compiled from Rieken 1988; Gaupp et al. 1993; Platt 1993, 1994; Zwingmann et al. 1998, 1999; Schmidt-Mumm and Wolfgramm 2002)

Fluid-rock interactions

of brines along fault structures can be related to different stages of basin evolution and locally occurred along with the migration of differently composed gases (see Krooss et al. this volume). For example, fluid inclusions in authigenic quartz hosted in Carboniferous rocks indicate early CH4-CO2 migration during increased subsidence in Late Permian times and later fluid migration during Mesozoic subsidence (Lüders et al. 2005). Late entrapment of fluid inclusions under low (nearly hydrostatic) pressure conditions can be related to stages of Cretaceous inversion. The migration and entrapment of non-hydrocarbon gases in fluid inclusions seems to be related to this event (Krooss et al this volume). During (local) uplift and exhumation, meteoric water eventually mixed with, or even replaced, basinal brines. The formation waters became less saline and more oxidising with increasing dilution by meteoric water. Influx of meteoric water is inferred for inverted Rotliegend reservoirs in the UK and areas in the Netherlands, where locally Permian deposits were subaerially exposed during Late Jurassic and Early Cretaceous times. Although the formation of kaolinite by meteoric water in the Rotliegend is controversial (Ziegler 2006), other case studies show clear evidence that the influx of meteoric water may cause dissolution of cements such as carbonate, sulfate and halite, leaching of feldspar, precipitation of kaolinite and formation of iron oxides (e.g., Burley 1984; Bjørlykke et al. 1989).

6.3.4 Organic-inorganic interactions Petroleum fluids generated during maturation of organicrich deposits (Gaupp et al. this volume) may interact with mineral constituents of sedimentary rocks and may alter their porosity and permeability characteristics significantly. In general, reactions may occur within source rocks, migration pathways and reservoir lithologies. Solid bitumen in veins or in the pore space of sediments, and oil inclusions in authigenic minerals, are direct evidence of the former presence of liquid petroleum. Solid bitumen occurs as extremely thin to massive impregnations on pore surfaces of porous rocks, especially on authigenic clay minerals in sandstones. However, evidence for organicinorganic interactions, coming mainly from diagenetic studies of reservoir rocks, is only indirect and difficult to decipher because in most cases only reaction products are preserved. The processes have to be inferred from the spatial distribution, sequence and chemistry of authigenic minerals, the availability of petroleum source rocks, the timing of organic maturation with respect to inorganic diagenesis and from experimental studies. In many cases, it remains difficult to prove whether organic-inorganic

reactions or other important driving forces, for example temperature and/or the chemical-mineralogical contrast between sandstones and interbedded mudstones, were responsible for the observed diagenetic patterns. The volumetrically dominant components capable of being expelled from source rocks are liquid hydrocarbons (oil), gaseous carbon dioxide (CO2), methane (CH4) and in some cases nitrogen in the form of ammonium (NH+4) and hydrogen sulfide (H2S). Products of early organic maturation include organic acids, but are generally not well characterised. In general, petroleum components in sedimentary rocks (i) affect or control redox reactions, for example with iron and sulfate, (ii) influence pH, which has an effect on feldspar, clay mineral and carbonate stability, (iii) influence the mobility of specific metal cations by complexation, and (iv) may reduce or stop the growth of minerals by physically coating the pore surfaces in oilwet systems. In particular organic acids, carbon dioxide, liquid hydrocarbons, ammonium and hydrogen sulfide are likely to interact with inorganic components in source and reservoir rocks. The reduction of dissolved sulfate by hydrocarbon compounds produces H2S, HCO3- and a solid bitumen residue (Machel et al. 1995b; Krooss et al. this volume). Bacterial sulfate reduction is thought to be restricted to temperatures below 80 °C in most cases, while thermochemical sulfate reduction appears to require temperatures above 140 °C (Worden et al. 1995), or 100-140 °C (Machel 2001). Dissolved sulfate is commonly derived from dissolution of solid sulfate minerals such as anhydrite or gypsum. Replacement of Ca sulfates by carbonates is a widespread consequence of bacterial or thermochemical sulfate reduction. If metal ions such as Fe are present, the generation of H2S results in the precipitation of sulfides, e.g., pyrite. Many red bed reservoirs such as the Rotliegend sandstones of the CEBS are affected by bleaching due to reduction of ferric iron coatings around detrital grains. The original presence of such coatings in grey sandstones is evident from iron oxide relics preserved beneath early cement overgrowths, and from the fact that the distribution of bleached zones is not facies-related. A large number of case studies (e.g., Kharaka et al. 1986; Surdam et al. 1989; Garden et al. 2001) as well as experimental work (Shebl and Surdam 1996) indicate that bleaching can be caused by liquid hydrocarbons or organic acids. Bleaching of red beds is often recorded in the vicinity of faults and fractures, which can be regarded as conduits for reducing petroleum fluids. H2S-bearing carbon dioxide fluids may also have the capacity to bleach red beds (Haszeldine et al. 2005; Palandri and Kharaka 2005). Although methane can reduce haematite, this reaction can only be expected

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R. Schöner · V. Lüders · R. Ondrak · R. Gaupp · P. Möller Figure 6.3.10. Simplified sequence of diagenetic processes and products in Rotliegend reservoirs of two contrasting margins of the North German Basin (modified from Gaupp et al. 1993; Schöner and Gaupp 2005)

in deeply buried sandstones at temperatures above approximately 160 °C (Barker and Takach 1992; Giggenbach 1997). Petrographic and geochemical data suggest that substantial volumes of iron can be removed from bleached sandstones by migration of reducing fluids (Chan et al. 2000; Haszeldine et al. 2005). On the other hand, iron mobilised during bleaching may also be stored in-situ by formation of Fe2+-rich sheet silicates or carbonates within bleached sandstones (Schöner and Gaupp 2005). Bleaching does not necessarily involve a significant change in the bulk iron content of the rock. Seepage of hydrocarbons from petroleum reservoirs may lead to reduction of ferric oxides/ hydroxides and formation of magnetic minerals, especially magnetite, in the overlying deposits (e.g., Donovan et al. 1979; Machel 1996). Associated magnetic anomalies have been used for petroleum exploration. Feldspar dissolution and clay mineral growth in sandstones is enhanced by acidic/CO2-rich pore fluids, although the migration of acidic fluids is controversial and many other factors control feldspar and clay mineral stability. For example, feldspar leaching may be caused

by meteoric water flushing at shallow burial depth (e.g., Bjørlykke 1983a, 1984), or may be thermodynamically controlled (e.g., Dutton and Land 1988; Wilkinson et al. 2001). Import of highly reactive organic acids and CO2 from source rocks has long been considered as the reason for major leaching events in sandstones (Schmidt and McDonald 1979; Surdam et al. 1984). However, the capacity of such fluids to cause significant leaching in sandstone reservoirs without direct contact to source rocks has been doubted since highly reactive components probably interact with minerals in the source rocks or along migration pathways before they enter reservoir horizons (Bjørlykke 1984; Giles and Marshall 1986). On the other hand, experimental studies indicate that substantial amounts of organic acids and CO2 may be generated within reservoirs by oxidation of n-alkanes (Seewald 2001, 2003). The controls on clay mineral formation in reservoir sandstones during burial diagenesis are very complex and a matter of controversial discussions since decades (Wilkinson and Haszeldine 2002; Ziegler 2006; and

Fluid-rock interactions

Figure 6.3.11. Schematic map of Pre-Upper Rotliegend geology in the CEBS, with superimposed Rotliegend sandstones at the southern and northern margin of the basin (modified from Lokhorst et al. 1998; Baldschuhn et al. 2001; Hoffmann et al. 2005; and others)

others). However, there is compelling evidence that migration of petroleum-bearing fluids is one important control. It has been suggested that pervasive fibrous or platy illite in reservoir sandstones precipitates from increasingly acidic/CO2-rich pore waters during early stages of oil charging (e.g., Cookenboo and Bustin 1999; Barclay and Worden 2000). Impregnation of illite surfaces by solid bitumen suggests that clay mineral growth terminated after oil migration or emplacement. Worden and Morad (2003) pointed out that clay reactions slow down in water-wet sandstones after oil immigration, and cease if the sandstones are oil-wet. Authigenic illites found in coal layers and organic rich shales are often characterised by high NH4+ contents. The principal source of nitrogen in sediments is the thermal or microbial decomposition of organic matter. NH3 released by decomposition of organic material may be incorporated in authigenic illites (Daniels and Altaner 1990) or migrate with pore fluids and oil (Stahl 1977; Williams et

al. 1989). Authigenic illites in primary nitrogen poor volcanic Permo-Carboniferous rocks of the NGB (<100 ppm NH4+-N) may reach up to 1000 ppm, a clear indication of migration of NH4-bearing fluids into these rocks (Mingram et al. 2005; Krooss et al. this volume). A comparative study of Rotliegend sandstones from the northern and southern part of the NGB indicates that organic-inorganic interactions cause a substantial rearrangement of diagenetic patterns in deeply buried red bed reservoirs (Schöner and Gaupp 2005; Fig. 6.3.10). Fluvial and aeolian sandstones at the southern margin of the Rotliegend depocentre are underlain by thick Carboniferous sediments containing abundant coal and minor oil source rocks (Fig. 6.3.11). In contrast, it is unlikely that significant amounts of organic maturation products migrated into the sandstone belt representing the northern margin of the basin in Rotliegend times. A number of diagenetic patterns were only observed in Rotliegend sandstones from the southern part of the basin, where a hydraulic

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Figure 6.3.12. Schematic model showing the effect of petroleum migration on diagenetic mineral reactions in the Rotliegend of the NGB as suggested by comparison of two contrasting basin margins (see also figure 6.3.11). Burial histories, temperatures and oil generation are examples taken from calibrated 1-D basin models (example for the southern basin margin based on data of Schwarzer and Littke 2007)

contact to the Carboniferous is evident or likely (Fig. 6.3.10): Bleaching, substantial dissolution of feldspar and early carbonate/sulfate cements, pervasive illite and local kaolinite/dickite growth, subsequent impregnation of pore surfaces by solid bitumen, precipitation of Fe2+-rich sheet silicates and carbonates and local extensive late quartz cementation. Illite authigenesis has been dated to about 180200 Ma in most areas (Gaupp et al. 1993; Zwingmann et al. 1998; Liewig and Clauer 2000). This age corresponds to the late phase of main oil generation in the underlying Carboniferous source rocks (Fig. 6.3.12). Authigenic kaolinite formed only within Carboniferous sandstones and around direct hydraulic contacts between source rocks and Rotliegend reservoirs (Gaupp et al. 1993). Feldspar leaching and clay mineral authigenesis are most intense close to fault zones enabling such hydraulic connections. This implies a direct influx of acidic fluids from the Carboniferous into the Rotliegend (Fig. 6.3.8). The spatial coincidence of these phenomena with the availability of maturing source rocks, plus the timing of diagenetic reactions and petroleum generation, suggest that petroleum fluids interacted with red bed reservoirs in multiple ways. The described characteristic diagenetic succession may be used to trace palaeo-petroleum migration pathways. Areas that were not affected by petroleum-bearing fluids, such as the Rotliegend at the northern margin of the NGB, apparently retained cementation patterns dominated by early to shallow burial diagenetic processes. Although the reactivity of acidic, CO2-rich fluids that were expelled from

the Carboniferous into the Rotliegend may have decreased with increasing distance from hydraulic contacts, oxidation of hydrocarbons by formation water or ferric oxides within the Rotliegend could have been an alternative source of organic acids and CO2. The co-occurrence of bleaching and mesodiagenetic leaching processes has also been observed by Muchez et al. (1992) and Surdam et al. (1993). Although oil migration was not post-dated by any significant late cementation in many cases, there are some examples in the Rotliegend of the CEBS, where large intergranular volumes were filled by late (post-bitumen) quartz cements. The close intergrowth of quartz and bitumen stained illite, and oil-inclusions within quartz cements, suggest a connection between this late quartz cementation and petroleum migration, but the processes are not yet understood.

6.3.5 Modelling fluid-rock interactions Modelling of fluid-rock interactions or more general geochemical reaction modelling is an approach to better understand mineral dissolution and precipitation processes in the pore space of rocks and to quantify the diagenetic evolution of sedimentary rocks. Natural systems are very complex and simplifications are necessary to illustrate the principal processes and parameters controlling the temporal evolution of the system. Fluid-rock interactions are generally very slow, covering geological time spans of thousands to

Fluid-rock interactions Figure 6.3.13. Gypsum solubility at 25 °C as a function of sodium chloride concentration calculated according to the Debye-Hückel theory (B-dot equation) and the Pitzer equations (modified after Kühn 2004). Full squares show data of Block and Waters (1968)

several hundreds of thousands of years. Geochemical models can simulate these processes in short time periods thus providing a convenient test of concepts and ideas. A properly constructed model is the simplified representation of the complex natural system which can be easily evaluated and applied to the problem in question (Bethke 1996). Fluid-rock interactions can be modelled with increasing degrees of complexity, from geochemical equilibrium through reaction path and kinetic models to coupled transport and reaction. The short description presented here just scratches the surface of this topic but extensive reviews of modelling water-fluid interactions in various environments have been published elsewhere (e.g., Bethke 1996; Giles 1997; Parkhurst and Appelo 1999; Clauser 2003; Kühn 2004; Steefel et al. 2005). The first and most critical step in setting up a geochemical model is conceptualising the system or process of interest in a useful manner (Bethke 1996). The system defines that part of nature which is relevant for the model. In a closed system, the chemical composition is fixed while in an open system transport of species into and out of the system is possible. Thermodynamic equilibrium systems take centre stage in any geochemical model. They contain no spatial or temporal information in their most simple formulation. The equilibrium system contains an aqueous fluid, different mineral species and perhaps a gas phase. If temperature, pressure and composition are known, the equilibrium state can be calculated. Thermodynamic equilibrium calculations assume ideal solutions characterised by vanishing small concentration of dissolved species in the fluid phase. Under these conditions, the activity of a species corresponds to its concentration in solution. In nature, concentrations of dissolved species are always

higher than the ideal conditions defined in thermodynamic equilibrium calculations. Species activities in non-ideal solutions have to be corrected for ionic interactions to obtain actual concentrations. The Debye-Hückel theory assumes that ions behave as charged spheres interacting by Coulomb forces (e.g., Bethke 1996). The Debye-Hückel equation is suitable to calculate activity coefficients (γi) in low to moderately high concentrated solutions (I < 0.1):

(6.3.11)

in which zi is the charge number of the ion, åi is an ion size parameter, A and B are temperature-dependent constants and I is ionic strength. The ionic strength can be calculated as I = ½ ∑ mi zi 2

(6.3.12)

in which mi is molality. The Davies (1962) equation is a variant of the DebyeHückel equation suitable for brackish water with ionic strength I > 0.1 to I ~ 0.5 (Stumm and Morgan 1996):

(6.3.13) Helgeson (1969) presented another activity model similar to the Davies equation: (6.3.14)

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Figure 6.3.14. Simple model for the relative change in feldspar composition of a Triassic sandstone from the Central Graben, North Sea. Decrease of K-feldspar and increase of albite is modelled by heating sandstone in contact with highly saline pore fluid under closed system conditions simulating increased burial

called “B-dot” equation, where B· is another temperature dependent parameter. The B-dot equation is parameterised up to 3 molal (mole/kg) ionic strength and 300 °C (Helgeson and Kirkham 1974). For solutions with higher ionic strength the ion interaction model or Pitzer equations (Pitzer 1973, 1975; Pitzer and Mayorage 1973, 1974; Pitzer and Kim 1974) are the most suitable approach. However, the present lack of data for aluminum species and the limited availability of data for higher temperatures still prevent its application in the realm of diagenetic processes which are strongly influenced by aluminum silicates like feldspars and clay minerals. Therefore, the different variations of the Debye-Hückel equation are most widely used in fluid-rock interaction models but the modeller should be aware of the errors in the mass balance calculations introduced by applying activity coefficients defined for dilute solutions to highly concentrated brines. The effect of increasing ionic strength and the application of different activity correction on mineral solubility can be demonstrated on gypsum solubility (Fig. 6.3.13) which increases significantly with increasing salinity. While all activity models produce similar results at low ionic strength of the solution the different activity models deviate significantly with increasing ionic strength indicating the limitations of the models and emphasising the necessity to calculate actual concentrations with the most appropriate activity correction equation. The geochemical equilibrium models can be used to calculate the ion speciation in the fluid and the saturation state of fluid with respect to the different mineral species for a given temperature, fluid composition and mineral assemblage. After speciation of the system, the resulting precipitation or dissolution reactions of minerals and the

resulting new concentrations of dissolved species can be calculated to obtain chemical equilibrium. After the initial condition of the studied system is calculated it is possible to calculate its reaction path as a function of temperature or compositional changes. Polythermal reaction models commonly model the evolution of a closed system without mass-change within the system as in studies of groundwater geothermometry or the interpretation of laboratory experiments. Fluids in hydrothermal systems, for example, are sampled at high temperatures but analysed at room temperatures. The in-situ conditions can be reconstructed by calculating the equilibrium state at room temperature and by modelling a polythermal path to estimate the fluid chemistry at reservoir conditions. Another simple example for a polythermal closed system reaction is albitisation of K-feldspar with increasing temperature. Petrographic data of Triassic reservoir sandstones from the Central Graben in the North Sea show that about 2 vol.-% of the K-feldspar is dissolved and about 0.5 vol.-% albite is formed (Lippmann 2007, pers. comm.). A possible explanation for the observed albitisation is heating of the highly saline pore water in the pore space providing sufficient sodium for albite precipitation. The process can be modelled by heating of a reference volume of sandstone under closed system conditions from 50 to 150 °C (Fig. 6.3.14). This indicates that albite formation does not require an import or export of species, while the dissolution of a four times larger amount K-feldspar requires open system conditions at least on a local scale. The simplest open models are titration models. A reactant is gradually added to a solution in equilibrium and dissolves, changing the original equilibrium. The process may cause minerals to become oversaturated and precipitate or already existing minerals to become undersaturated and dissolve. Titration models may be used to model fluid mixing with the reactants being a second fluid, or to simulate evaporation in which the reactant, water, is removed. However, in most cases a mineral or a gas phase like carbon dioxide is added to the system and the resulting dissolution and precipitation reactions are calculated. Such models are used to simulate the interaction of a rock with a pore fluid. The reaction of very small amounts (in the order of 10-3 moles) of a silicate mineral is commonly sufficient to saturate a fluid with respect to the mineral. The considerable amounts of diagenetic cements commonly observed in sedimentary rocks require high water-rock ratios indicating that pore fluids must have been replaced many times to precipitate the observed amounts of cement. Minerals react at different rates: some react fast enough to maintain equilibrium, others react so slowly relative to the time period of interest that the reaction can be ignored, while the rest requires a kinetic description. Thermodynamic calculations only deal with the equilibrium state of

Fluid-rock interactions

Figure 6.3.15. Modelled alteration of a Triassic sandstone-mudstone succession from the Central Graben (North Sea). Transport of solutes occurred by diffusion only. Different mineralogical compositions and the resulting chemical gradients act as driving mechanism for transport and diagenetic alteration in particular at the contact of the different lithologies

a geochemical system. However, reactants and products do not achieve equilibrium instantaneously but approach it with time, especially in the temperature range realised in sedimentary basins. Therefore, kinetic processes should be taken into account, which led to the development of kinetic reaction models. For a detailed introduction the reader is referred to the comprehensive description of kinetic theory by Lasaga (1998). Coupled reaction-transport models combine geochemical modelling with transport processes which bring reactants into the system or remove them. Such models allow modelling of the diagenetic evolution as an integrated process of mass and heat transport and the resulting mineral reactions. Transport can be driven by fluid flow or diffusion, or both (Urai et al. this volume). Since mineral dissolution or precipitation changes porosity and permeability which control flow rates, a direct feed back of diagenetic processes on transport occurs. The quantification of transport processes coupled with geochemical reaction processes can help to decipher the most likely combination of mechanisms which can explain observed diagenetic alterations. For example, chemical gradients of neighbouring rocks of different (mineral) composition such as shale and sandstone, together with diffusive transport, can be sufficient on geological time scales to drive diagenetic alterations. This was shown by simulations of Thyne et al. (2001) for K-feldspar dissolution and illitisation in a shale-sandstone succession of the Late Triassic to Early Jurassic Statford Formation in the Viking Graben of the

North Sea. In an alternating shale-sandstone sequence of Triassic reservoirs from the Central North Sea albitisation of K-feldspar in the sandstone and illitisation of smectite in the shales was observed (Fig. 6.3.15; Lippmann 2007, pers. comm.). In addition, the intensity of quartz cementation in the sandstone increases towards the sand-shale contact. The developed conceptual model assumes a diffusion driven export of potassium from the sandstone into the shales. The illitisation of smectites releases sodium and magnesium which are exported into the sandstones where sodium assists albitisation of K-feldspar. A simple model scenario calculated with CrunchFlow (Steefel et al. 2005) shows that the geochemical gradient and diffusional transport are sufficient to reproduce the observed albitisation of K-feldspar and the increased quartz precipitation of the sandstones close to the shale-sand contact (Fig. 6.3.14). The observed illitisation of smectites is also reproduced by this model. Modelling fluid-rock interaction provides not only results but also uncertainty about the accuracy of those results. In fact, uncertainty is an integral part of modelling that deserves as much attention as any other part of the study (Bethke 1996). Many input parameters are poorly constrained. Although the data base is growing continuously, thermodynamic data are not available for all minerals found in nature. In particular, the complex composition of clay minerals poses great difficulties when modelling the diagenetic evolution of clastic rocks. The calculation of activities in highly concentrated saline solution cannot be performed accurately without a complete data set of (experimentally determined) parameters for the Pitzer

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model, which is currently not available. Kinetic data for minerals is even more incomplete and inconsistent, although recent efforts help to reduce the gap (Palandri and Kharaka 2004). The determination of reactive mineral surface areas is still very difficult and a major source of uncertainty (Fischer and Gaupp 2005). These problems have to be considered when modelling results are interpreted. A sensitivity study with methodic variation of input data which is not well constrained will narrow likely scenarios and provide error ranges.

6.3.6 Geological applications Sedimentary basins contain the bulk of known natural resources. Apart from coals and syngenetic sedimentary ore deposits, the formation of these natural resources is bound to the available pore spaces or fractures, and the fluids within these pores or fractures. Hot saline waters itself may be used for medical purposes and thermal energy. Sediments containing non-saline freshwater, which is confined to relatively shallow aquifers in most cases, are very important reservoirs for drinking water. However, a complete review of all aspects of fluid-rock interactions with respect to natural resources is far beyond the scope of this chapter, which just highlights some important geological applications in the CEBS. Fluid-rock interactions are responsible for various kinds of mineral deposits in the CEBS (Walther and Dill 1995), but economically important deposits also occur in sedimentary strata of adjacent areas. An important type of mineral deposit that has been mined over centuries in Germany and Poland is the Kupferschiefer at the base of the Zechstein. The mineralisation, which is dominated by copper, lead and zinc ores, is at least partly related to diagenetic metal enrichment, although a primary syndepositional concentration of metals is likely (Bechtel et al. 2001). Economic carbonate-hosted massive lead-zinc mineralisation in Ireland deposited during early diagenesis of their Carboniferous host rocks (e.g., Wilkinson 2003). Mississippi-Valley-type Pb-Zn-(F-Ba) deposits hosted by Palaeozoic to Jurassic rocks occur in Upper Silesia (Poland) and the Pennine Orefields in England (e.g., Dunham 1983; Leach et al. 1996). Besides water aquifers and mineral deposits, hydrocarbons are very important resources of sedimentary basins. In the CEBS, commercial oil reservoirs were recovered predominantly in Late Triassic, Jurassic, Cretaceous and Palaeogene strata of northern Germany and the North Sea (e.g., Boik 1981; Glennie 1998). Locally, commercial oil reservoirs also occur in Devonian sandstones and in Zechstein carbonates. Natural gas exploration mainly

concentrated on the Upper Rotliegend sandstones that were deposited along the southern margin of the CEBS. Other important gas reservoirs were discovered in Carboniferous (mainly Namurian to Stephanian) sandstones, in Zechstein carbonates and in the Triassic Buntsandstein (e.g., Lokhorst et al. 1998). Jurassic sandstones are the most important gas reservoirs in the central and northern North Sea, but Cretaceous and Palaeogene sediments also host commercial gas deposits (Glennie 1998). Because of the great economic importance and the extensive data set available, the relevance of fluid-rock interactions to reservoir properties of the Rotliegend will be discussed here in more detail. For economic applications, the spatial-temporal evolution of petrophysical attributes like porosity and permeability must be reconstructed. Ideal reservoir rocks have porosities greater than 10 vol.-% and permeablilities exceeding 1 mD. Diagenetic processes which destroy pore space by mineral precipitation or enhance porosity by mineral dissolution, therefore, are of great interest for exploration. As discussed previously, pore fluids act as transport media for reactants which drive mineral reactions. The major controls on reservoir properties are sedimentary facies, burial history and maximum burial depth, mineralogical composition and diagenesis. Sedimentary facies controls the initial porosity and permeability due to grain size, sorting, matrix content and sedimentary texture. Considering the continental Rotliegend sediments, well sorted dry aeolian dune and sandflat environments as well as lake shore environments commonly comprise sandstones with the best initial reservoir properties. Damp or wet aeolian environments are much less favourable owing to their detrital clay content that reduces both permeability and porosity. Fluvial sandstones have highly variable primary porosities and permeabilities, depending on their compositional and textural maturity and matrix content. Although depositional environment is the first control, subsequent mechanical compaction and mineral precipitation/dissolution are the dominant processes controlling reservoir properties at depth (Fig. 6.3.16). Generally, porosity and permeability decrease with depth due to compaction and cementation. In the absence of early framework stabilising cements, mechanical compaction decreases the intergranular volume of well sorted sandstones from about 40-45% at surface to around 26% at 2500 m depth, as long as the components are rigid (Paxton et al. 2002; see also Urai et al. this volume, box 6.2.1). The intergranular volume equals porosity only if cements and matrix are absent and can be described as maximum potential porosity. Sandstones rich in mechanically unstable lithic rock fragments such as shales or volcanic grains are affected by more intense mechanical compaction that may result in intergranular

Fluid-rock interactions Figure 6.3.16. Porosity-permeability-plot of dryaeolian Rotliegend sandstones (North German Basin, data from Gaupp et al. 2005). This graph illustrates the wide spread of permeabilities for a given porosity, even if depositional facies is kept constant. All theses sampled sandstones had an optimum depositional facies favouring the potential of best reservoir qualities. Both porosity and permeability are significantly impeded by diagenetic alteration (see text) in samples towards the lower left edge of the cluster

volumes of less than 10% due to ductile grain deformation. Sandstones with abundant clay coatings around detrital grains tend to become strongly compacted, since the presence of K-bearing sheet silicates at grain-grain contacts favours quartz dissolution during compaction (“pressure solution”) (Houseknecht 1988; Bjørkum 1996). Diagenesis can reduce, destroy, and in some cases enhance porosity and permeability. In areas with elevated evaporation, much of the pore space of good reservoir lithologies may be occluded by blocky carbonate, sulfate or chloride cements shortly after deposition. Owing to selective early cementation, rocks with less ideal depositional porosity/permeability characteristics may then remain as preferential flow paths for compaction fluids. The growth of radial, grain-coating chlorite inhibits quartz cementation, and hence has the potential to preserve porosity in advanced burial diagenesis (e.g., Ehrenberg 1993). Chlorite-cemented sandstones are among the best Rotliegend reservoir rocks in the CEBS (Gaupp et al. 1993, 2005). Where the grain framework is stabilised by mechanical compaction and only minor cementation, sandstones may still have reasonable reservoir quality. Examples are Rotliegend red beds comprising haematite grain coatings together with low volumes of quartz overgrowths on detrital grains. Burial diagenetic enhancement of porosity and permeability may take place by dissolution of grains, especially feldspar, and early cements creating so-called secondary porosity. Since parts of the pore space generated by dissolution may be occluded concurrently or subsequently by further cementation or by compaction, the timing of dissolution is crucial for reservoir quality. A major period of dissolution in Rotliegend sandstones of the CEBS can be observed prior to illite authigenesis and oil migration (Fig. 6.3.10 and 6.3.12). However, this leads to reservoir quality improvement only in places where illite

growth was negligible or minor. The permeability of sandstones can be reduced by more than two orders of magnitude due to growth of authigenic clay minerals, especially due to platy (meshwork) and fibrous (hairy) illite, which increase tortuosity, specific inner surface areas and tend to plug pore throats (“tight gas sandstones”). Therefore, platy and fibrous illite in deeply buried reservoir sandstones is a major risk for deep gas exploration, even if the porosities are relatively high (e.g., Almon and Davies 1981; Leveille et al. 1997). Petrographic observations in Rotliegend reservoirs also imply a local post-illite, porosity-generating dissolution of feldspar and cements, as well as a pervasive local pore-plugging, bitumen-stained quartz cementation (Gaupp et al. 1993, 2005; Schöner and Gaupp 2005). Both late-stage processes are still poorly understood, but can be essential for exploration. Late-stage dissolution can lead to an inversion of reservoir properties from former poor (cemented) reservoirs to moderate or good reservoirs. In summary, reservoir properties of deeply buried sandstones are a complex function of depositional environment, compaction, cementation and mineral dissolution, with the depositional characteristics becoming increasingly less important at depths below about 3000 m. The diagenetic patterns are heterogeneous, cross-cut facies boundaries, and are difficult to predict. The understanding of these patterns requires a fundamental understanding of the fluid evolution and of fluid-rock interactions in reservoir lithologies within the context of the geological history of the basin. Another aspect is the storage potential of sedimentary basins for methane, CO2, or other fluids/gases. The commonly fluid-bearing pore space of subsurface geological formations is of increasing economic and ecological importance concerning recent and future energy and envi-

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ronmental problems. In the past, sedimentary basins were mainly regarded as potential hosts of hydrocarbon reservoirs in the underground. However, because of the growing acceptance of the relevance of green house gas emissions on global climate change, its usage for gas storage, especially for sequestration of CO2, the most abundant greenhouse gas, is of special interest. A combination of oil production and CO2 injection into the reservoir is applied by American oil companies since the 1970’s to optimise well production rates of depleted oil/gas repositories (e.g., Jessen et al. 2005). Whereas these projects originally focused on enhanced oil/gas recovery (EOR/EGR), natural gas production accompanied by CO2 sequestration scientific studies started in 1996 in the Sleipner gas field, Norwegian North Sea (Saline Aquifer CO2 Storage – SACS; e.g., Kongsjorden et al. 1998; Torp and Gale 2004). These investigations and many other recent studies attempt to verify the effects of technically induced CO2 into deep saline aquifers or exploited hydrocarbon reservoirs (e.g., Baines and Worden 2004; Brosse et al. 2005; Förster et al. 2006; and references therein). CO2 injection can result in mineral dissolution processes (e.g., carbonate, chlorite, feldspar) and thus improve porosity and permeability conditions for CO2 storage. Precipitation of carbonate minerals, for example, calcite, dolomite, siderite and dawsonite, is one of several concepts for permanent storage of carbon dioxide in the pore space of sedimentary rocks (e.g., Gunter et al. 2004). On the other hand, mineral precipitation may, in the worst case, occlude the pore space and impede further CO2 injection. Although exploited gas reservoirs are considered suitable for the permanent storage of CO2 since they have accumulated gas for many millions of years,

carbon dioxide might react with minerals in the seal which could lead to very slow leakage into higher strata above the reservoir horizon. The fundamental understanding of potential CO2-(pore)fluid-rock reaction processes and the careful monitoring during and after CO2 injection is therefore essential for all CO2 sequestration projects. Whereas these technologies focus on minimising industrial greenhouse gas emissions, the geothermal heat of sedimentary basins may be used for non to low emission energy production on a local to regional scale. Geothermal energy production in sedimentary basins is based on the injection of relatively cold water into deep porous or (artificially) fractured units and its controlled ascent to surface as heated fluids. The best potential hot water horizons in the CEBS are the sandstone units of the Rotliegend, Buntsandstein and Keuper (compare Stollhofen et al. this volume). However, thermal and hydrological disturbances brought about by deep water injection may initiate mineral dissolution and precipitation processes, at worst resulting in cementation of the reservoir (e.g., Pape et al. 2005). In conclusion, the pore space and the fluids within sedimentary rocks are of great economic and environmental importance. The consequences of just how this resource is used must be carefully considered and conflicting interests must be reconciled. For example, the artificially induced mineral reactions described above which are positive for CO2 sequestration might be a problem for possible later geothermal energy production because the mineral precipitation destroys the pore space and hydraulic conductivity.

6

Chapter 6.4

Petroleum systems R. di Primio · B. Cramer · C. Zwach · B.M. Krooss · R. Littke

6.4.1 Concepts of petroleum system modelling The evolution of sedimentary basins is the result of a complex combination of physical, chemical and microbial processes. In order to study such processes various methods are applied in academia and industry. Most of these methods can be divided into three major groups: 1. Observations, such as well sample analysis or analysis of seismic data. 2. Experiments are used to study processes under controlled laboratory conditions. The controlled heating of source rock samples in a laboratory is used for example to gain information on the maturation of organic material and the resulting generation of hydrocarbons as a function of time and temperature. 3. Numerical modelling may be applied to simulate and quantify physical and chemical processes using a computer. Numerical models can be as simple as a porosity-effective stress function for a given lithotype, and may be as advanced as three-dimensional multiphase reservoir simulators. The first two groups, observation and experiment have been used to study geological phenomena for several hundred years. The latter group, numerical modelling, is a relatively young discipline and will be the focus of this chapter. All three groups together provide the general base of our understanding of a natural system. This set of methods is also used in many other disciplines such as physics and biology. However, each one of the three groups has certain shortcomings regarding the study of a sedimentary basin. Observations only give “snap-shot” information of the geological history and are usually restricted to a few wells and outcrops. Laboratory experiments, when dealing with geological processes, often suffer from the much shorter time-scale in which they are performed; identified relationships such as kinetic parameters then have to be extrapolated over several orders of magnitude and may carry a large error. Numerical models tend to oversimplify the multitude of physical and chemical processes occurring

in sedimentary basins. However, in combination, these three methods provide a powerful toolbox for the analysis and quantitative description of basin evolution, especially with respect to petroleum system analysis. Magoon (1988), and later Magoon and Dow (1994), coined the term petroleum system. In their landmark publications, a petroleum system was defined as a natural system that encompasses a pod of active source rock and all related oil and gas, and which includes all the geological elements and processes that are essential if a hydrocarbon accumulation is to exist. The essential elements of a petroleum system include the source rock, reservoir rock, cap rock and overburden. The main processes are trap formation and hydrocarbon generation, migration and accumulation. The characterisation of a petroleum system requires the geochemical measurement and mapping of source rocks and associated products, whereas the description of the timing of the relevant events requires the geological reconstruction of basin evolution. Magoon and Dow (1994) proposed a set of four figures (a burial history to determine the time of the critical moment where the source expelled hydrocarbons, a map of the accumulations at the time of the critical moment, a cross section of the petroleum system at the critical moment and an events chart to summarise the petroleum system), a table (listing all accumulations in the system) and a name as the best way to sum up the geographic, stratigraphic and temporal evolution of a petroleum system. The limitation of the definition of a petroleum system to include only a single source rock formation and a single reservoir formation soon became watered down, especially after it was recognised that many petroleum accumulations represent mixtures of hydrocarbons derived from different sources at different maturities (Wilhelms and Larter 2004). Today it appears more common to speak of petroleum systems (plural) active in a basin whereby the burial history and the events chart are the two figures which have become the standard output to describe petroleum system evolution. Figure 6.4.1 shows an example of a petroleum system event chart, in this case a chart available for download from the website of the USGS that has been modified for the Central European Basin in order to describe the most important source rock, reservoir rock and seal units.

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Figure 6.4.1. Mesozoic petroleum system events chart for the onshore part of the Central European Basin System (modified from USGS website and Kockel et al. 1994). The Palaeozoic natural gas system events chart is shown in Bayer et al. (this volume)

From the definition of a petroleum system it becomes evident that a modelling approach which reconstructs the processes and events taking place as a function of geological time (i.e., puts everything into a temporal framework) is the best way to investigate and quantify the evolution of a system. Petroleum system modelling was the tool which developed in response to these needs. In this text, “petroleum system modelling” implies the numerical simulation of physical and chemical processes in sedimentary basins and its application to describe the petroleum system evolution (reservoir, non-reservoir and source rocks). Petroleum system modelling has been shown in the past to be a very effective learning tool for exploration geologists and is essentially aimed at predicting the occurrence of hydrocarbon accumulations in the subsurface. As is the case with most simulation techniques, petroleum system modelling implies a simplified description of complex and dynamic physical and chemical processes over geological time. It always simplifies as much as possible and acceptable to mathematically simulate the results of natural processes relevant for petroleum system analysis within practical computing times (hours to days). Table 6.4.1 gives an overview of the disciplines and information types involved in basin modelling. In many programs natural processes are quantified by forward modelling, i.e., by simulating a process from the

geological past to the present. As an example, porosity evolution of a reservoir rock during burial is simulated from an initial porosity at deposition and its subsequent reduction as a function of mechanical and chemical compaction. These programs, hence, integrate mathematical descriptions of the most relevant individual physical and chemical processes and are used to analyse the complex, non-linear behaviour of natural systems until the present state is reached. Petroleum system modelling has been in use for at least 20 years. A variety of publications, including textbooks, exists in which the methodologies used are described in detail. The purpose of this chapter is not to re-iterate common knowledge but to focus on advances made in the last few years. For an overview of the fundamentals of petroleum system modelling, the interested reader is referred to the textbooks by Tissot and Welte (1984), Giles (1997), Welte et al. (1997) and Allen and Allen (1998). Advances made in the last few years have mainly been in process understanding. Specifically, as a reaction to the change in hydrocarbon exploration focus from mainly oil to oil and gas the entire concept of the “oil window” is under re-evaluation. Oil stability at high pressures and temperatures is a critical issue: up to which temperatures can we expect to find liquid-rich hydrocarbons (i.e., not only oil but also condensate-rich gases)? What are the lowest temperatures for the generation of thermogenic hydrocar-

Petroleum systems

Table 6.4.1. Overview of important disciplines and types of information required in petroleum system modelling Geology

Geophysics

Geochemistry

sequence stratigraphy

geometries

source rock quality

sedimentary processes

seismic facies

maturation

reservoir geology

seismic attributes

timing of generation and expulsion

structural geology

direct hydrocarbon indicators (DHIs)

petroleum composition

heat flow history

petrophysical properties

preservation; biodegradation

bons (oil and/or gas)? Is petroleum alteration predictable? Is it possible to predict the hydrocarbon composition generated by a source rock? How do source rock depositional environments control the ensuing petroleum properties? These questions will be addressed in the following chapters, generally in the context of the Central European Basin system.

6.4.2 Petroleum Source Rocks A viable exploration programme not only requires the presence of traps identified by seismic analysis, but also an analysis of petroleum charge. The latter depends on the presence of source rocks, their thickness, distribution, organic richness and quality (Katz and Pratt 1993). Petroleum source rocks are generally siliciclastic and carbonate rocks rich in organic matter, as well as coals. The deposition of source rocks is generally restricted to marine and lacustrine subaquatic environments in which organic matter is deposited faster than it can be destroyed, but can also include peat swamps and mires as precursors of coal (Tourtelot 1979; Littke et al. 1997). One characteristic feature of depositional environments in which petroleum source rocks are formed is the low amount of oxygen present. Precursors of source rocks are often sapropels, in which oxygen was completely absent at the time of deposition. Similarly, oxygen contents are generally low in peats, grading to zero within the uppermost centimetres of the sediments. This is to a great extent due to the fact that the abundance of organic matter leads to a rapid consumption of oxygen due to the activity of aerobic bacteria consuming organic matter. Nevertheless, not all organic matter rich sediments are deposited under conditions of complete oxygen deficiency. Two prominent examples of recent depositional environments for sapropel (source rock) formation are shown

in figure 6.4.2, together with indications of the oxygen content. In silled oceanic basins (such as the Black Sea) or silled lakes (such as Lake Tanganyika) with limited water circulation, in which moderate to high bioproductivity exists in surface waters, stratification of the water, e.g., due to superposition of low salinity water derived from rivers over saline, marine waters, can lead to anoxic conditions in the deep water (Fig. 6.4.2, upper part). Thereby, the further decay of organic matter is inhibited, leading to better conservation. In contrast, there are also areas where a stronger water circulation exists, leading to the presence of many nutrients in surface water. In such a case very high bioproductivity may result (such as offshore Peru or SW Africa). Oxygen concentrations under such conditions are strongly diminished but not zero (Fig. 6.4.2, lower part), i.e., the suboxic-anoxic boundary is at shallow depth in the sediments, but not in the water. Below the zone of aerobic degradation of plant organic matter, anaerobic degradation occurs. In shallow marine environments, sulphate reducing bacteria are the main drivers of this decay. In contrast, most fresh water and peat environments contain much less sulphate than sea water; therefore degradation by methanogenic bacteria is of greater importance there. Degradation via these processes leads to the loss of a large percentage of the former plant material, especially of the more labile parts. This holds true even for environments with a high preservation potential. Geochemists commonly divide the resulting organic matter into two groups: kerogen and bitumen. Kerogen is the insoluble (in organic solvents such as dichloromethane) organic matter and forms the bulk of the total sedimentary organic material. Bitumen is soluble and includes all conventional oil and gas. Kerogen classification is of the utmost importance in petroleum charge studies since abundance, quality and maturity will have the greatest impact on the presence

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R. di Primio · B. Cramer · C. Zwach · B.M. Krooss · R. Littke Figure 6.4.2. Generalised models for the deposition of organic matter-rich sediments. Upper part: stagnation model for anoxic silled basins; lower part: productivity model for upwelling regions (modified after Littke and Welte 1992)

of petroleum reservoirs. Kerogen quantity is commonly measured by analysis of the total organic carbon content (TOC), which is a method commonly applied in many geological projects. The most widely applied method to study kerogen quality is the Rock-Eval pyrolysis method described in detail by Espitalié et al. (1977) and Tissot and Welte (1984). The derived Generation Potential (S2), Hydrogen Index (HI) and Oxygen Index (OI) values are commonly used to qualify kerogen as Type I, II (both being oil-prone), III (coaly, gas-prone) or IV (inert, no generation potential). A typical kerogen classification for some major source rocks in the Central European Basin System (CEBS) is shown in figure 6.4.3. Plotted are data from the western Lower Saxony Basin. For detailed analyses of petroleum systems, Rock-Eval pyrolysis should not be used as the only method to describe kerogen quality. For example, gas/oil ratios generated from kerogen with similar HI-OI values can differ greatly. Also, the heterogeneity of the kerogen is not represented in the simple parameters derived from Rock-Eval, and mixtures of variable particles cannot be detected thereby. Therefore additional methods such as pyrolysis coupled with gas chromatography (Horsfield 1997) or microscopy (Taylor et al. 1998) are highly recommended. An example for the characterisation of sediments by organic petrological methods is shown in figure 6.4.4.

Kerogen quantity and quality are combined in SPI (source potential index) values. SPI is a measure of the maximum quantity of petroleum that can be generated within a vertical column of source rock under 1 m2 of surface area. Combined with knowledge of the drainage area, this information can significantly help to quantify petroleum charge. However, realistic estimates require in addition information on the timing of petroleum generation in the context of kerogen maturation. Kerogen maturation depends greatly on the temperature evolution of source rocks and leads to petroleum generation. A great number of maturation parameters have been developed in order to quantify maturity; some of them are described in Littke et al. (this volume). Among them, vitrinite reflectance and Rock-Eval Tmax are probably the most widely applied ones. It should be stressed, however, that no maturity parameter exactly mimics bulk oil generation or gas generation, since the kinetics of the respective processes are different. An ideal petroleum system model should, therefore, include source rock kinetics measured on representative samples from the respective basin. In reality, this often proves to be difficult because no representative, unweathered, immature samples are available. More general background on petroleum source rocks, organic facies, and organic geochemistry of source rocks can be found in Jones

Petroleum systems

Figure 6.4.3. Plot of Rock-Eval S2 (Generation Potential) versus Total Organic Carbon (TOC) for three of the most important source rock intervals in the southern (onshore) part of the CEBS. Low S2-values of some Toarcian samples are due to high thermal maturity. Data from Adriasola-Muñoz et al. (2007)

(1987), Huc (1988), Katz and Pratt (1993), Littke (1993), Littke et al. (1997) and Killops and Killops (2005). In the Central European Basin System, different types of source rocks exist at various stratigraphic levels. These include type I lacustrine kerogen from the Wealden formation, marine type I-II kerogen in the early Toarcian Posidonia Shale and in Late Jurassic shales of the North Sea, carbonate-evaporite type II-III kerogen in the Late Permian Zechstein formation and coaly type III source rocks in the Late Carboniferous interval. Thus, the CEBS has basically the full source rock/kerogen type inventory known from literature, making it highly complex in terms of petroleum systems. The youngest source rocks are of Early Cretaceous and Late Jurassic age. During this time interval, shallow marine and partly terrestrial depositional environments predominated, leading to deposition of source and reservoir rocks. Some Early Cretaceous intervals contain abundant organic matter, but are generally immature (Littke et al. 1998). An exception are the Berriasian Wealden Shales which may be mature at favourable positions in the basins and contain locally, especially in the western part, very hydrogen rich kerogen (Fig. 6.4.3). There, mean TOC values are 7% and the thickness is at about 25 m. Laterally, Wealden Shales grade eastwards into less oil-prone terrestrial sediments containing coaly organic matter (Wiesner 1983; Kockel et al. 1994).

In the North Sea, Late Jurassic source rocks are very prominent. The Kimmeridge Clay of Kimmeridgian to Ryanzian age is possibly one of the best studied source rocks in the world. It is characterised by excellent hydrocarbon potential and TOC contents reaching up to 10% (Barnard and Cooper 1983; Cornford 1998). The Kimmeridge clay is usually subdivided into a “hot” and “cold” unit based on the gamma ray response, whereby the hot unit is characterised by higher kerogen quality and quantity. Various lithostratigraphic schemes are used to define the LateJurassic in the North Sea. In the UK and East Shetland basins the Kimmeridge Clay extends generally from the base of the Kimmeridgian to the end of the Ryazian and is underlain by the Oxfordian Heather formation, which is generally characterised as containing a poorer terrestrially influenced kerogen type. In the Danish Central Graben the same two lithostratigraphic units are called the Farsund and Lola formations, whereas in the Norwegian Central Graben the Mandal (hot shale) and Farsund (cold shale) formations cover the Kimmeridgian to Ryazian interval and are underlain by the poorer quality Haugesund formation. In the Viking Graben, the Kimmeridge Clay interval is called the Draupne formation while the Oxfordian source unit is also called the Heather formation (Fraser et al. 2003). In Haltenbanken, the Spekk formation covers the Oxfordian to Ryazian interval and is underlain by the Callovian Melke formation. In the Barents Sea only the Hekkingen forma-

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R. di Primio · B. Cramer · C. Zwach · B.M. Krooss · R. Littke Figure 6.4.4. Petrographic composition of organic matter deposited in small oceanic basins (rifted continental margins, back arc basins) and along passive continental margins (upper part) as well as in progradational submarine fans, central oceans, on guyots (eroded volcanic islands), and in trench-slope transition zones (lower part) of the deep sea (after Littke and Sachsenhofer 1994)

tion of Kimmeridgian to lower Volgian age contains significant source potential. In addition to the Late Jurassic, source potential is also encountered in Middle and Early Jurassic coals (e.g., Callovian Pentland and Bryne formations of the Central and Viking Grabens, Hettangian to Sinemurian Åre formation in Haltenbanken; Husmo et al. 2003), as well as to a minor extent in Triassic and Permian sequences and in individual Cretaceous horizons. One of the best studied source rock intervals worldwide is the early Toarcian Posidonia Shale that was deposited in

almost the entire area of the CEBS and further towards the south and west. If immature, it contains a very oil-prone kerogen with HI values in the order of 700 (Fig. 6.4.3; Rullkötter et al. 1988; Adriasola-Muñoz et al. 2007). The Posidonia Shale is early mature to mature in large parts of the CEBS, but reaches very high levels of maturation (overmature stage) at the southern margin (southern Lower Saxony Basin). Its depositional history and petroleum generation characteristics have been summarised in numerous publications (e.g., Littke et al. 1988, 1991;) and it has to be regarded as the most important oil source rock of

Petroleum systems

Box 6.4.1 Prospect risking In petroleum exploration different prospects have to be ranked depending on their chances of success with respect to finding an economically viable amount of petroleum. Such risking is important in order to be able to provide input for the economic evaluation of the prospect and to rank prospects against each other. Prospect risking commonly involves a detailed assessment of the individual elements of an effective petroleum system. These include trap (type, presence, timing with respect to HC-generation, retention time, seal properties, pore pressure), reservoir (presence, quality) and source (quality, volume, migration, preservation). Usually prospect risk is expressed as a percentage of the probability of discovery, where 100% indicates that a discovery is guaranteed. An example of a prospect risking scheme is shown in figure 1. For each element to be risked an individual value representing its likelihood is calculated. For example the source rock is assessed based on: a proven, mapped, continuous source rock layer would be assigned a probability of 100% (or 1) · Presence: is a thick enough source present? Assuming seismic visualisation is not good enough and well data indicate highly · Volume: variable source rock thickness, this factor would be of a lower probability, say 50% (0.5) based on vitrinite reflectance data the source is early to mid mature, i.e., not optimal, hence also a probability of · Maturity: 50% (0.5) Assuming these three factors are sufficient for addressing source rock risk, the individual risks or probabilities are multiplied delivering a total source rock probability. In our fictitious case Presence (1) x Volume (0.5) x Maturity (0.5) = Source probability 0.25 (or 25 %) That means that we expect a 25% or 1 in 4 chance of finding the optimal source rock conditions. Also the probabilities of the individual elements are multiplied by each other to arrive at a final prospect risk. Generally this type of assessment is not performed in a deterministic way. Petroleum companies generally try to incorporate uncertainty into the numbers used to describe probabilities. Accordingly, statistical methods are a very important part of prospect risking. Details on this subject and on prospect risking in general can be found in Harbaugh et al. (1995). In all aspects of prospect risking hydrocarbon volumes, phase and quality represent key issues. Expected hydrocarbon volumes control the economics of a prospect. In addition phase state, i.e., gas or oil or a two-phase system, can significantly alter the economics (e.g., if you have oil production facilities near your prospect you do not want to find large amounts of gas). Petroleum quality can also change the economics of a given prospect, as for example biodegraded heavy oil is of lower economic value than unbiodegraded light oil. Biodegradation is usually a risk at reservoir temperatures below 80 °C. However, recent research indicates that uplifted reservoirs which have been subjected to temperatures higher than 80 °C have a low risk for biodegradation (Wilhelms et al. 2001). The latest developments with respect to hydrocarbon preservation evaluation include, accordingly, a burial history dependent assessment of biodegradation risk. Eventually, risking of biodegradation will also have to include models with higher time resolution in order to estimate biodegradation rates (Larter et al. 2003). Figure 1 Key elements of a risking scheme

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R. di Primio · B. Cramer · C. Zwach · B.M. Krooss · R. Littke Migration pathways change dynamically through geological history, due to complex geometric changes as a result of, e.g, tilting, juxtaposition, folding, etc. Only a full 3D modelling of source, reservoir and carrier evolution can be assumed to be close to the real petroleum system evolution. 2D migration models can, however, illustrate general migration mechanisms. An alternative to 3D basin modelling (assessing volumes generated as compared to trap volumes) is 2 ½ D modelling (e.g., ray tracing) or advanced gridding of 1 and 2D basin modelling results to reach a near-3D representation (Zwach et al. 2000). Additional points to consider during prospect risking:

· · · ·

when is a source rock not a risk? thick, rich, oil window mature occurrence proven (well data) maturation after trap formation and cap consolidation when is migration not a risk? reservoir in direct contact with the generative source proven carrier system proven fill-spill trend when is volume not a risk? generated volume orders of magnitude higher than trap pore volume when is preservation not a risk? high charge rate in the recent past

· ·· · ·· · ·

Every company uses a different scheme! But all schemes include most of the elements shown in figure 1. In addition to the absolute risk estimated for a given prospect it should also be taken into account that a prospect with a large volume and higher risk might be economically more interesting than a low volume - low risk prospect.

the CEBS (with the exception of the northwestern North Sea part). This source rock typically has a thickness of 30 to 60 m and the mean TOC is at an immature stage close to 10%. In some parts of the basin, e.g., in the north (Schleswig-Holstein, Rodon and Littke 2005) and west (Netherlands, Nelskamp et al. 2007), it was partly eroded during Late Jurassic/Cretaceous inversion. The largest oil field of the onshore part of the CEBS, the Mittelplate field, was sourced by Posidonia Shale (Grassmann et al. 2005). Source rock potential of the thick Triassic and Permian sequence is rather limited. An exception is the Late Permian Zechstein evaporite formation which contains an organic matter rich carbonate layer as basinal facies (Stassfurt carbonate-Ca2). This carbonate layer grades laterally into high energy, shallow water oolitic carbonates which can act as reservoir rocks for oil and gas. Thicknesses are at about 10 m and average TOC values below 0.5%, but locally higher than 1% (Hindenberg 1999). In the central eastern part of the CEBS (east German area), this source rock is partly in the oil generation stage, whereas it is much more mature in the west, where it contributes to gas generation and accumulation (Littke et al. 1996). Zechstein-derived oils are isotopically heavier than Mesozoic oils (Wehner 1997). Similar intra-salt source rock sequences grading laterally into reservoir rocks are also known from other parts of the world, e.g., from the Cambrian Ara Formation in Oman (Schoenherr et al. 2007a). The major gas source rock in the CEBS is the Late Carboniferous coal-bearing sequence which completely underlies the Permian-Mesozoic basin except for the

northernmost- northeastern part. The entire sequence reaches a thickness of more than 3000 m in the south, where more than 100 coal seams occur. Towards the north and northeast, this sequence becomes thinner and fewer coals occur. In the south, where the Carboniferous occurs close to the surface allowing for coal mining, up to 5% of the sequence consists of coal. In addition there is abundant dispersed organic matter in silt- and sandstones, which is mainly coaly in nature. Accordingly, Hydrogen Index values are only moderate (up to 300) for the coals. In the CEBS, large gas accumulations such as Groningen have been fed by this coal-bearing sequence. Maturity ranges from the early gas generation or even oil generation stage in the offshore area between Netherlands and Germany or in the Ruhr basin in the south to highly overmature in the southern Lower Saxony Basin and in the area of Hamburg/southern Schleswig-Holstein. In general, it is expected that the deep-lying Carboniferous coals produce gas at high levels of maturity. A further contribution to this thermal gas from deeper source rocks of Early Carboniferous and Devonian age is probable, but these formations are not well explored due to their great depth. However, shallow gas generation also takes place in the CEBS, e.g., in uplifted coal sequences at the southern margin.

6.4.3 Shallow and microbial gas The occurrence of hydrocarbon gas at relatively shallow depths is of interest both for economic and drilling

Petroleum systems

security reasons. Shallow gas is commonly observed in seismic surveys, leading to a brightening of reflectors at levels where high saturations occur. The limits to which gas accumulations can be termed shallow are arbitrary; in this volume we will set the limit to 1000 m depth. The source of the accumulated gas can be both microbial or thermogenic, and it can be derived both locally or be the product of hydrocarbon migration. Understanding the processes leading to shallow gas formation and accumulation is of paramount importance in reducing drilling risks and enhancing exploration possibilities. In the autumn of 2005, the offshore exploration well 35/2-1, Norwegian North Sea drilled by Hydro at 384 m water depth, encountered a commercial gas accumulation at a depth of less than 200 m below mudline. The Peon discovery is, thus, probably one of the shallowest commercial gas accumulations worldwide, and has additionally opened a new play type in the Norwegian North Sea. Shallow gas accumulations have up to date been perceived more as a drilling hazard than as a possible opportunity. The Peon discovery has, however, highlighted the economic potential of such shallow gas occurrences.

Some shallow gas systems are characterised by the possible origin of gas via microbial rather than thermogenic processes. Previously, the depth ranges in which microbial and thermogenic processes occurred were deemed to be separate, whereby a temperature boundary of roughly 80 °C was assumed to mark the end of microbial reactions and the initiation of thermogenic ones. Today this view is changing. Conventional assessment of the onset of thermogenic hydrocarbon generation is usually based on predictions from source rock kinetic models. These characterise the generation of hydrocarbons from source rocks assuming a distribution of activation energies and a single frequency factor, measured experimentally, as descriptive of the reactions taking place both in the lab and in nature. Dieckmann (2005) developed a new method of processing experimental hydrocarbon generation data to provide coupled activation energy-frequency factor information. For terrestrially influenced, heterogeneous source rock types this processing led to a shift of 30-40 °C in the predicted onset of hydrocarbon generation, displacing the first generation of hydrocarbons to temperatures as low as 60 °C under natural heating conditions. Such predictions are actually also supported by observations of high gas contents in rocks of low maturities (Snowdon 2001).

In the greater Mackenzie River Delta area of northwestern Canada large amounts of gas are trapped in the form of comparatively shallow gas hydrate accumulations. In 1998 the Mallik 2L-38 gas hydrate research well was drilled in this area in order to investigate the occurrence of gas hydrates (Dallimore et al. 1999). The results reported indicate the occurrence of up to 10 significant gas hydrate-bearing units within the depth interval from 810 to 1102 m. estimated that the Mallik gas hydrate accumulation may contain up to 110 billion m3 of gas, indicating the potential of gas hydrate accumulations as an energy resource. Gas hydrate accumulations have been reported from various marine and permafrost regions in a depth range between 130 and 2000 m (Collett 2002). The estimated amount of gas stored in hydrate accumulations worldwide may exceed the volume of known conventional gas resources (Kvenvolden 1988), although newer assessments indicate that these numbers may possibly need significant revision (Milkov 2004). Among the factors controlling gas hydrate occurrence, the availability of gas ranks paramount.

The bulk kinetic characterisations of hydrocarbon generation cannot distinguish the hydrocarbon type being generated, i.e., oil or gas. In general, the observation that at low maturities source rocks generate liquid-rich fluids whereas with increasing maturity gas contents increase is largely valid. Muscio and co-workers (1994), however, documented that in specific source rock types, in this case the Bakken Shale (Williston Basin, USA), large proportions of gaseous hydrocarbons are monitored in samples of low maturity (down to 0.3% vitrinite reflectance). There, gas content actually decreases with increasing maturation and is attributed to the aromatic nature of the immature source rock kerogen and/or to a specific biological precursors (e.g., diaromatic carotenoid structures from green sulphur bacteria). Another example is western Siberia, where some of the largest gas fields on earth exist at relatively shallow depths of 800 to 1500 m. These shallow gas accumulations have a composition and an isotopic signature which suggest an origin from very early (low temperature) thermogenic reactions (Cramer et al. 1999; Littke et al. 1999).

Besides gas hydrates, a variety of relatively novel unconventional shallow gas systems have received a large amount of interest in recent years. Coalbed gas (see Ayers 2002 for a summary) is among the better known and well explored unconventional gas systems, whereby the latest developments have extended exploration targets from pure thermogenic gas-sourced systems to also include microbial coalbed gas.

As the onset of thermogenic processes moves towards shallower conditions in sedimentary basins, microbial activity has been recognised in increasingly deeper and hotter settings. In the past decade the temperature limit for life has been extended to values as high as 113 °C (Blöchl et al. 1997) or even 121 °C (Kashefi and Lovley 2003). Despite these advances in recognising temperature boundaries for life, the limits of the subsurface biosphere

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Box 6.4.2 Hydrocarbon Generation Kinetics The generation of hydrocarbons from organic rich sediments proceeds via a multitude of parallel and successive, quasi-irreversible chemical reactions which are controlled by chemical kinetics. The progress of these reactions follows the Arrhenius equation k = A e(-E/RT) where k is the reaction rate, A the pre-exponential factor, E the activation energy, R the gas constant and T the absolute temperature. Theoretically each reaction occurring during kerogen (solid sedimentary organic matter) breakdown has its own solution for the Arrhenius equation. In order to simplify the calculation procedure it is commonly assumed that a single, average preexponential factor and a distribution of activation energies (either Gaussian or discrete) is sufficient to correctly describe the hydrocarbon generation process. Hydrocarbon generation kinetics are determined experimentally, and kinetic laws allow the mathematical link between high temperature – short time conditions in the lab and low temperature – long time conditions of sedimentary basins, thus permitting the extrapolation of laboratory to natural heating rates. Artificial maturation experiments are performed on immature source rock or kerogen samples using open or closed system conditions and isothermal or non-isothermal conditions. Open system experiments exclusively address the primary reactions of hydrocarbon generation, i.e., the direct conversion of organic matter to hydrocarbons. Closed system experiments allow, in addition, the characterisation of secondary reactions among the primary products. The non-isothermal open system experimental approach is currently the most common, with non-isothermal closed system methods being used for compositional characterisation and investigation of thermal petroleum alteration either in the source rock (cracking of residual oil) or in the reservoir (cracking of accumulated oil). Kinetic models of bulk petroleum formation (bulk kinetics) address only the timing and amount of petroleum generated. Compositional kinetic models provide, in addition, information on the timing and amount of individual components or component groups of petroleum, i.e., gas and oil generation or generation of compound classes such as wet gas (C2-C5), light hydrocarbons (C6-C14) and long chain hydrocarbons (C15+). Such compositional kinetic models incorporating primary and secondary cracking characterisations can be integrated with phase models for the calculation of petroleum properties (GOR, API gravity, phase state) in the subsurface. Bulk kinetics of different types of source rock show a large variation, which can be attributed to differences in the molecular structure of the parent kerogen. The molecular structure depends on the types and amounts of biological precursor material preserved during organic diagenesis and additionally on the diagenetic incorporation of sulphur into the kerogen matrix (Tegelaar and Noble 1994). Examples of the variability of bulk kinetics among different kerogen types are shown in figure 1. Type I kerogens usually show high petroleum potential (proportion of the organic matter which can be converted to hydrocarbons) and a very narrow activation energy distribution, indicating a very homogeneous organic matter type (dominance of a single macromolecular constituent derived from precursor organisms). Type II kerogens show lower potentials and wider activation energy distributions and are typical for marine source rocks. Type II-S kerogens show in part enhanced petroleum potentials as compared to the typical Type II, but their activation energy distributions tend to be wider and shifted towards lower values than those of sulphur-lean kerogens. Finally, Type III kerogens show the lowest petroleum potential and also a relatively broad activation energy distribution (Fig. 1; see also Gaupp et al. this volume). Figure 1. Examples of pseudohistograms of activation energy distributions of different kerogen types

Petroleum systems

In basin modelling, selecting the appropriate kinetics representative of the source rock studied can be a problem. Often samples of the source rock are not available for analysis. There are several solutions to this problem: so called “typical” bulk kinetics have been established in the literature, representing single kinetic measurements of hydrocarbon generation of well-known source rocks (e.g., the Late Jurassic Kimmeridge Clay). Alternatively, models have been published which allow the prediction of kinetic properties based on source rock sedimentary facies and age (Pepper and Corvi 1995). When using laboratory defined kinetics the question remains whether the extrapolation of the kinetic data determined on immature samples derived (usually) from basin margin settings is valid. The success of such a prediction depends strongly on how closely the samples resemble the actual source rocks generating petroleum in the deep basinal settings. An alternative to the extrapolation of kinetic data determined on source rocks is the use of analyses of petroleum asphaltenes. Asphaltene kinetics have been demonstrated to reflect the extent of thermal stress to which the source rocks were subjected when the petroleum was generated and hence, in combination with basin modelling, offer information on where and when the petroleum analysed was generated (di Primio et al. 2000). In-reservoir oil to gas cracking has been the topic of much discussion in the last decade. Various laboratory analyses of petroleum and individual hydrocarbon stability pointed towards a relatively high stability of liquid hydrocarbons even under geological heating conditions. The threshold temperatures required to initiate in-reservoir oil-to-gas cracking were consistently predicted to be higher than 170 °C. The cracking of residual oil to gas in source rocks has, however, been observed to occur under markedly lower temperatures, possibly due to the catalysing effect of high molecular organic materials such as residual kerogen. Accordingly, it is likely that liquid hydrocarbons are thermally more stable in a reservoir than in a source rock, and that the ultimate oil vs. gas distribution in a sedimentary basin is mainly controlled by primary and secondary reactions of hydrocarbon generation and degradation in the source rock.

with respect to microbial methane generation or microbial degradation of petroleum seems to remain at a temperature between 65 and 90 °C (Head et al. 2003), whereby sediments which have been exposed to higher temperatures seem likely to be devoid of microbial populations (archaea and bacteria) capable of degrading hydrocarbons (Wilhelms et al. 2001). The distinct separation of microbial and thermal processes is, in view of the above, no longer clear cut within sedimentary basins. Boundaries of microbial and thermal processes tend to overlap to a sometimes significant extent. Horsfield and co-workers (Horsfield et al. 2006) discuss the possibility of microbial activity being dependent on thermal cracking processes. The authors demonstrated that the rates of thermal degradation of organic matter (to microbially amenable compounds) closely match rates of respiration and electron donor consumption in deep biosphere environments. This observation leads to the possibility that abiotically driven degradation reactions can actually provide the substrates necessary for microbial activity in deep sediments. If this observation is true, then kinetic expressions of these abiotic reactions, readily measured in the lab, could be applied to predict feedstock generation for deep microbial systems, i.e., addressing the microbial gas generation window. Natural hydrocarbon seepage is a process recognised to occur in varying intensity along most continental shelves. The seafloor and seismic expression of such seepage sites includes gas chimneys as well as surface and buried mud

volcanoes (Milkov 2000), pock marks (Hovland and Judd 1988) and carbonate mounds (Hovland and Thomsen 1997). Using basin modelling, Naeth et al. (2005) provided strong indications that the locations of carbonate mounds observed in the deep water of the Porcupine Basin, Ireland, correlated to sites of preferred hydrocarbon leakage. In the Orange Basin of South Africa seismic mapping of gas chimneys has been used to identify locations of gas leakage, and their occurrence linked to the active petroleum system (Paton et al. 2007). Often gas leakage is due to seal failure as the result of high gas columns overcoming the capillary entry pressure of the cap rock pores (Urai et al. this volume). Major consequences of hydrocarbon seepage include the support of diverse microbial communities as discussed above, the potential to reduce the stability of continental margin sediments and, importantly, the flux of carbon into the hydrosphere and atmosphere (Judd et al. 1997). The proportion of methane reaching the atmosphere due to natural seepage has until recently been assumed to be insignificant (IPCC 1990). However, first assessments of the flux of gas from seabed seepages indicate that the geological sources of atmospheric methane may be more significant than is generally acknowledged. Judd et al. (1997) reported that estimates of gas seepages on the UK continental shelf can reach values representing 2 to 40% of the total UK methane emission. In Germany, coal mining along the southern margin of the CEBS releases high amounts of methane, turning Germany from a methane sink into a methane source area. A significant part of this gas is of shallow, possibly microbial origin, as revealed by isotope data (Thielemann et al. 2001). Ac-

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cordingly, a deeper understanding of shallow methane gas exchange with the hydrosphere and atmosphere is necessary to obtain more quantitative data on the global carbon cycle and to monitor sources of the 14C-free part of atmospheric methane, this being an important greenhouse gas.

6.4.4 Sources of deep gas Deep and tight gas accumulations have currently become major exploration targets in the petroleum industry both worldwide and in the CEBS. They belong to the important unconventional fossil energy resources in sedimentary basins, together with coalbed methane, gas hydrates, tar sands and oil shales. Tight gas plays are gas accumulations hosted in low-permeability reservoir rocks. The unconventional gas system receiving the highest amount of interest in the last five years is shale gas. The term shale gas is derived from the fact that gas is produced from tight shales, which act as source, reservoir and seal of the gas. Gas in these reservoirs is typically, but not exclusively, of thermogenic origin. Thermogenic hydrocarbons are generated and expelled during the thermal maturation of organic matter in source rocks. Petroleum generated directly from kerogen is the result of so-called primary cracking, whereas that generated by the thermal degradation of primary products is the result of secondary cracking. In both cases gaseous and liquid hydrocarbons are produced which, depending on the pressure and temperature conditions in the subsurface, occur either as a single phase or as two phases (see Box 6.4.3). The proportion of gas in the fluids generated increases as a function of increasing thermal stress, becoming the dominant product during secondary cracking (di Primio and Skeie 2004). The main source of deep gas is assumed to be the secondary cracking of unexpelled hydrocarbons in source rocks (including coals). While the primary cracking of kerogen to petroleum has been characterised in detail (Béhar et al. 1992, 1997; Burnham 1989; di Primio and Horsfield 2006; Schenk and Horsfield 1997) the mechanisms of secondary cracking still require investigation. For example, the stability of liquid petroleum in reservoirs seems to be relatively high (Horsfield et al. 1992; Schenk and Horsfield 1997), indicating that it can withstand temperatures close to 200 °C under geological heating rates, whereas in source rocks residual oil is converted to gas at lower levels of thermal stress (Dieckmann et al. 1998). Accordingly, the kinetics of the secondary cracking reactions lead to 30 °C differences or more in the predicted onset of secondary cracking under natural heating conditions, ranging from 150 °C (in-source cracking) to well over

180 °C (in-reservoir cracking) at geological heating rates. Accordingly, compositional predictions from kinetic models in combination with basin modelling have encountered great difficulties in correctly predicting the fluid composition and phase in deep, hot reservoirs (Vandenbroucke et al. 1999). However, newer models calibrated to observed fluid properties indicate that good predictions regarding the gas to oil ratio (GOR) and even phase properties are possible for the realm of primary cracking (di Primio and Horsfield 2006; di Primio and Skeie 2004). The characterisation of secondary cracking mechanisms and products generated is usually done using closed system pyrolysis experiments (Béhar et al. 1992; Dieckmann et al. 1998; Schenk et al. 1997). In these experiments kerogen is transformed to bitumen and petroleum which then, at higher temperatures, crack to increasingly lighter compounds. Latest experimental results indicate that recombination processes occurring within source rocks during maturation may result in the formation of additional, unexpected gas potential which leads to the generation of predominantly methane at maturities exceeding 2.5% VRr (vitrinite reflectance) (Dieckmann et al. 2006; Erdmann and Horsfield 2006). The mechanisms leading to the generation of this late gas involve a recondensation process where first-formed high molecular weight compounds in the closed system yield a macromolecular material that undergoes secondary cracking at elevated temperatures. The high temperature secondary gas formation from cracking of the recombination residue as detected for the terrestrially influenced and/or deltaic source rocks is quantitatively important and is thought to occur at very high temperatures (temperature of maximum conversion approx. 250 °C) for geological heating rates. The prediction of a significant charge of dry gas from such sources at very high maturity levels has important implications for petroleum exploration in ultra-deep basins where conventional kinetic models predict the source rocks to be essentially cooked out. The compositional evolution of petroleum in nature during increasing thermal stress is characterised by an increase in the proportion of methane. Compositional information from natural maturity series of genetically related petroleum fluids (light oils to gas condensates) demonstrates this relationship as shown in figure 6.4.5. Fluids with a GOR over 600 Sm3/Sm3 already belong to the gas condensate class and can be assumed to contain a significant proportion of products generated by secondary cracking. With respect to the compositional evolution of petroleum fluids, a comparison of the GOR evolution between natural petroleum fluids and closed system pyrolysis experiments performed on the corresponding source rock

Petroleum systems

Box 6.4.3 PVT Modelling PVT analysis and PVT modelling applications in petroleum production and exploration commonly refer to the reconstruction of petroleum fluid behaviour and physical properties at varying Pressure, Volume and/or Temperature. Both approaches have been routinely used by reservoir engineers to characterise the physical properties of a reservoired fluid as well as the changes in volume and phase state occurring during production (McCain 1990; Pedersen et al. 1989). PVT analysis involves the reconstruction of the reservoir fluid composition at reservoir pressure and temperature using analytical techniques such as gas chromatography, distillation, and density and molecular weight measurements. The samples required for such analyses are obtained either from the separator or from down-hole samples. In both cases gas and liquid fractions have to be first separated, analysed individually and using different techniques, and the resulting compositions recombined mathematically, using the volumetrics of the two phases observed, to obtain the reservoir fluid composition. In PVT modelling, reservoir fluid descriptions can be input into software packages which calculate PVT and phase behaviour based on fluid composition using equations of state (EOS). These EOS provide a mathematical description of the fluid behaviour for reservoir simulation and reserve estimation. In the last few years basin modelling simulators have been developed which allow taking fluid phase behaviour into account when modelling hydrocarbon flow from the source rock to the reservoir. The phase behaviour of a fluid is commonly described using either specific physical properties of the fluids such as formation volume factor (Bo), saturation pressure (Psat) or gas to oil ratio (GOR), or by means of a phase envelope plotted on a pressure vs. temperature diagram, as shown for typical fluid types in figure 1. The five fluid types shown in this figure represent the typical fluid terminology used by reservoir engineers. It should be taken into account that from the point of view of petroleum geochemistry all fluid types shown in figure 1 can be generated from a single source rock during increasing levels of maturity (di Primio et al. 1998). The shapes of petroleum phase envelopes depend on the fluid composition. The amount and composition of gas in the fluid strongly influence its saturation pressure and GOR, whereby a methane-rich gas composition results commonly in a higher saturation pressure than a comparably “wetter” gas composition. Amount and composition of the liquid fraction have a strong influence on the extent of the phase envelope to high temperature as well as on the critical point. Additionally, they indirectly influence the GOR and saturation pressure. Fluids with high liquid contents and high average liquid molecular weights and densities are characterised by low GORs and saturation pressures. The evolution of fluid composition during source rock maturation is reflected in the physical properties of reservoir fluids. Reservoirs which have been charged by the same source rock and differ only as a function of petroleum maturity show very systematic trends. Bo and Psat have, for example been seen to follow a linear trend for genetically related fluids (di Primio 2002). Thus, databases of fluid PVT properties can be used as a screening tool for oil population characterisation, relative maturity estimation, volumetric evaluation of complex reservoir filling histories and recognition of mixing effects. Latest developments in compositional kinetic modelling now allow the description of evolving petroleum fluid composition as expelled from the source rock, which is calibrated to petroleum phase behaviour (di Primio and Horsfield 2006). Figure 1. Phase envelopes of typical petroleum fluids: each phase envelope consists of a region where fluids occur in a single phase state and a region where they exist as two separate phases. The latter one is enclosed by a bubble point curve (green) and a dew point curve (red). The bubble point curve marks the PT-conditions where separation of a gas phase from a supercritical liquid phase takes place, while the dew point curve is defined as the PT-area where separation of a liquid phase from a supercritical gas phase occurs. The critical point (grey), located where bubble point and dew point curves meet, characterises fluid conditions intermediate between those of a liquid and a vapour phase. The line between reservoir conditions and surface conditions represents the PT reduction path the fluids go through during production

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R. di Primio · B. Cramer · C. Zwach · B.M. Krooss · R. Littke Figure 6.4.5. Increasing gas dryness (proportion of methane in the gas phase) as a function of increasing maturity (represented here by the Gas to Oil Ratio GOR) for a series of genetically related petroleum fluids

sample indicates that the use of a closed system configuration (i.e., kerogen and generated products remain in contact during the entire experiment) matches the observed natural GOR span best (di Primio and Skeie 2004), even extending far into the secondary cracking range. This has implications for the use of closed system pyrolysis experiments for the prediction of the GOR of fluids generated from a given source rock type. Di Primio and Horsfield (2006) demonstrated that GOR predictions from closed system pyrolysis experiments performed on source rock samples actually very closely match natural fluid GORs.

All types of pyrolysis experiments are, however, known to predict gas compositions incorrectly (di Primio and Skeie 2004; Mango 2001; Mango et al. 1994), which implies that for the correct compositional reconstruction of petroleum based on laboratory experiments gas compositions require correction. The physical properties of petroleum fluids change with increasing maturity. Fluid density decreases during increasing maturation and proceeding secondary cracking, such that the liquid phase API gravity increases from typiFigure 6.4.6. Evolution of saturation pressure with increasing maturity, here exemplified using the fluid GOR

Petroleum systems

cal unaltered oil gravities (25-40 °API) to those of light oils (40-45°API) and then condensates (>45 °API). This shift in density is associated with the decreasing proportion of high molecular weight compounds in the liquid phase. The saturation pressure of the evolving fluids also changes systematically. Box 6.4.3 shows the evolution of petroleum phase envelopes as a function of increasing maturation, shifting from a typical low saturation pressure (Psat) – high temperature shape to that of a high Psat – intermediate temperature to again low Psat – low temperature shapes. In view of the fact that the geological PT conditions represented by the straight line in the diagram do not vary much, we can conclude that genetically related fluids of increasing maturity should be characterised by low saturation pressures at low maturity, high saturation pressures at increasing maturities and again decreasing saturation pressures at the highest maturities. This is simply due to the fact that with increasing maturity the fluids phase envelope shifts to lower temperatures, eventually no longer being intersected by the geological PT line (the fluid does not reach saturation, i.e., wet gas). This behaviour is also observed in large fluid datasets, such as shown in figure 6.4.6.

6.4.5 Petroleum alteration biodegradation Microbial transformation processes in petroleum reservoirs, commonly denoted as biodegradation, may have an important impact on the quality, producibility and economic value of crude oils. Tissot and Welte (1984) state that among the world‘s total petroleum reserves, the heavy, degraded oils are as abundant as conventional crude oils. Head et al. (2003) even consider the world‘s reserves to be dominated by biodegraded heavy and super-heavy oils which occur as huge tar sand accumulations, especially in the foreland basins of North and South America. Biodegradation has become an important issue since exploration started to extend towards deep-water prospects with shallow burial depths and low geothermal gradients (Head et al. 2003). While the consequences of in-reservoir petroleum biodegradation are well known, the understanding and quantitative description of the actual processes involved remain vague. It is commonly believed that the prerequisites are

Figure 6.4.7. Biodegradation scale after Peters and Moldowan (1993). Shaded areas represent compound classes that are depleted or absent in the residual petroleum. The 25-norhopanes are formed at level 6 and destroyed at level 8

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Figure 6.4.8. Biodegradation scheme proposed by Wenger et al. Removal of selected compound groups at various levels of biodegradation (modified after Wenger et al. 2002 and Head et al. 2003)

the presence of (living) micro-organisms, substrate (electron donors) and electron acceptors such as sulfate. In addition, nutrients such as nitrate and phosphate are required as well. The concepts of biodegradation are closely linked to the notion of the “deep biosphere” which has emerged

only over the past decade. There is increasing evidence for the presence of micro-organisms in deep sedimentary systems inherited from the time of deposition. Infiltration of micro-organisms by flowing (meteoric) water is also conceivable. However, the observation that petroleum res-

Petroleum systems

ervoirs remain sterile after having been heated to temperatures (80-90 °C) that will kill all in situ petroleum degraders (Wilhelms et al. 2001) has led to the conclusion that micro organisms responsible for natural biodegradation processes in petroleum reservoirs are mostly indigenous. Biodegradation affects the molecular composition and the physical properties of petroleum in a characteristic manner. The preferential attack of n-alkanes results in an early removal of these compounds and in an enrichment of the residual oil in aromatic and polar (NSO) species. The associated reduction in the H/C ratio results in a higher specific density and thus decreasing API gravity values. The depletion in hydrocarbon species also causes a relative increase of trace metal contents, especially nickel and vanadium. Furthermore, biodegradation tends to increase the number of carboxylic functions and, thus, the acidity of the petroleum leading to an increase in the total acid number. The observation that this effect is less pronounced in carbonate reservoirs indicates the influence of the reservoir mineralogy on biodegradation (Wenger et al. 2002). The residual compounds in biodegraded petroleum are generally enriched in the 13C isotope (Connan 1984). Changes in the carbon isotopic composition of petroleum during biodegradation are mainly observed for n-alkanes with fewer than 18 carbon atoms (Sun et al. 2005). Also, within the aromatics, resins and asphaltenes, a trend towards 13C enrichment was observed. The enrichment of Table 6.4.2. Mechanisms of overpressure generation and their relative importance (Osborne and Swarbrick 1997) Overpressure Mechanisms Chemical compaction Disequilibrium compaction Tectonic activity Aquathermal pressuring Volume expansion: Smectite conversion Disequilibrium compaction: Smectite conversion Mineral diagenesis without dehydration Disequilibrium compaction Kerogen maturation Gas generation Differential density effect Potentiometric head Osmosis

Importance

·· ·· · ·

· ·· · ·

heavier isotopes during biodegradation in reservoirs can be described by the Rayleigh fractionation model and thus may be used for quantification of the process (Sun et al. 2005; Vieth and Wilkes 2006). As short chain n-alkanes are degraded preferentially and/or at a higher relative rate, their biodegradation increases the relative methane content and at the same time reduces the potential of the remaining oil to dissolve gas so that the gas-oil ratio is reduced (Wenger et al. 2002). Peters and Moldowan (1993) introduced a classification scheme which is now widely used for the characterisation of biodegraded oils (Fig. 6.4.7). The scheme is based on the observation that different petroleum compound groups and biomarkers are sequentially removed by microbial attack (Connan 1984) and are absent in the residual oil. The biodegradation rank is assigned a value from 1 (light biodegradation) to 10 (severe biodegradation). The 25-norhopanes play a special role in this biodegradation scale: they are formed at a biodegradation level of 6 and destroyed together with the bulk of the hopanes at a biodegradation level of 8. The Peters and Moldowan scale describes the biodegradation levels on a geochemical basis, irrespective of the relative masses of oil transformed and the quality reduction (decrease in API gravity value and increase in viscosity). From an economic point of view, the lower biodegradation levels where the most important oil quality reduction occurs are of major concern. Wenger et al. (2002) thus introduced a new scheme (Fig. 6.4.8) that provides a more detailed specification of the Peters and Moldowan biodegradation levels 1-5 in terms of individual compounds while lumping together levels 6-10 into one “severe” degradation class. Petroleum compositions deviating from the common biodegradation patterns, as exemplified by gas chromatograms showing light hydrocarbons co-occurring with a pronounced hump of “unresolved complex mixture” (UCM), are frequently observed. These observations indicate the practical limitations of the biodegradation scales which are, however, still useful for a first qualitative assessment of biodegradation levels. In most situations biodegradation will take place simultaneously with filling of the reservoir. According to present understanding, both processes occur at similar time scales (Larter et al. 2003). Depending on the relative rates of biodegradation and reservoir filling, the effects of biodegradation become less evident or may even be wiped out completely. In the CEBS, biodegradation of oils is common, ranging from moderate (Blumenstein et al. 2007) to severe (Horsfield et al. 1991). Gas alteration by thermochemical sulfate reduction (TSR) occurs as well (see Krooss et al. this volume).

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6.4.6 Overpressured reservoirs In deep parts of sedimentary basins, pore pressures are often significantly greater than hydrostatic pressures, the latter being calculated from depth, acceleration of gravity and average water density. This significant increase in fluid pressure, locally reaching almost lithospheric pressures, leads to changes in rates and directions of fluid flow, and also to great financial and safety risks in drilling operations. Therefore, prediction of overpressures before and during drilling has become a major challenge in earth sciences. Table 6.4.2 summarises some important aspects of overpressure generation; further information is gathered in Huffman and Bowers (2002) and Osborne and Swarbrick (1997). The table illustrates that undercompaction of mudstones due to rapid burial on the one hand and gas generation combined with kerogen conversion (solid-fluid conversion) on the other are the most important processes (Broichhausen et al. 2005). For overpressure generation, an additional requirement is fluid retention, i.e., the presence of sufficiently low permeability sediments. High pressure – high temperature (HPHT) reservoirs are increasingly becoming the focus of petroleum exploration

in the search for additional reserves. The processes leading to the accumulation and preservation of hydrocarbons in these settings are still poorly constrained, especially with respect to the occurrence of significant porosity at great depths and temperatures exceeding 150 °C. The more or less common occurrence of severely undersaturated light oils or wet gas-condensates is one peculiarity of HPHT reservoirs in the deeper grabens of the North Sea. Based on kinetic models these hydrocarbon compositions are virtually impossible to explain, as kinetic predictions indicate that essentially only dry gas should be generated from the source rocks under these conditions (e.g., Dieckmann et al. 1998; Wendebourg 2000). It seems likely that the problems encountered with hydrocarbon phase prediction in HPHT settings are due to either incorrect compositional kinetics, incorrect models of hydrocarbon expulsion or incorrect assumptions on the role of pressure on hydrocarbon stability. Di Primio and Neumann (2007) studied an HPHT setting in the Central Graben of the UK. There, a significant pressure increase is observed going through the Cretaceous Tor and Hod formations (Fig. 6.4.9). Within the Figure 6.4.9. Overpressure evolution in the UK Central Graben

Petroleum systems

that reservoir temperature alone was most probably not the controlling factor with respect to fluid composition. Using 3D basin modelling and including phase predictive compositional kinetic models calibrated to regional fluid properties plus a description of secondary cracking based on a combination of in-source and in-reservoir kinetics, the authors were able to reproduce the observed natural fluid compositions. The results indicated that source rock maturity and ongoing secondary cracking of unexpelled oil into gas exerted the main control on reservoir fluid composition. In addition, the authors demonstrated that gas generation through secondary cracking must have provided a significant additional source of pressure in the basin (see Table 6.4.2). The Haltenbanken area is also characterised by the occurrence of overpressured and normally pressured reservoirs which are mainly of Jurassic age. The boundary between hydrostatic and overpressured areas is generally fault controlled, with a single fault defining the boundary west of Åsgard field and a series of faults defining the more diffuse pressure transition further south (Fig. 6.4.10) (Karlsen et al. 2005) .

Figure 6.4.10. Overpressure distribution in the Haltenbanken area. Boundary between overpressured and hydrostatic area based on data from Karlsen et al. (2005). Map courtesy of the Norwegian Petroleum Directorate

Triassic reservoirs studied pore pressures reached values exceeding 800 bars and temperatures above 190 °C. Nevertheless, the reservoir fluids encountered in these settings were characterised by comparatively high liquid contents, with GORs spannning the range from 900 to 2600 Sm3/Sm3. Interestingly, the lower GOR was found in the deeper and hotter portion of the reservoir, indicating

Until recently the western part of the Haltenbanken was commonly known as a “dry, overpressured region” but this changed as Saga Petroleum drilled the Kristin prospect in a Jurassic play. This highly overpressured reservoir (overpressure exceed 400 bars) exhibits abnormally good porosity and permeability; porosities over 20% are reported at 5 km depth (Rønnevik 2000). Blystad et al. (1995) have described the structural setting in the region, typically dominated by the more shallow Halten terrace and deep Jurassic-Cretaceous basins in the Vøring Basin. Evidence from fluid inclusion analysis indicates that the main pressure boundary only developed in the course of the last 5-3 Ma (Karlsen et al. 2005) as a function of quartz cementation, the rate of which increases exponentially with increasing temperature, and occurs preferentially at fault zones. Figure 6.4.11. Comparison of modelled overpressure evolution in the overpressured Kristin reservoir, assuming disequilibrium compaction and the combination of compaction and gas generation

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Figure 6.4.12. Lithostatic and hydrostatic pressure trends and measured pore pressures in carbonate stringers of the infra-Cambrian Ara salt, Oman (after Schoenherr et al. 2007a)

One of the fundamental simplifications in current basin modelling packages is that sediment compaction is controlled by mechanical forces only. Accordingly, effective stress (overburden - pore pressure) is modelled to drive porosity reduction, with the result that overpressured rocks are modelled to have elevated porosity (undercompaction). The generation of overpressure in basin models is, hence, controlled by the location of low permeability barriers (cap rocks or faults) as well as by overpressure generation due to porosity reduction by mechanical compaction. However, in the overpressured parts of Haltenbanken this type of simulation results in a very slow increase in overpressure over a very long time. A rapid increase in overpressure during a short period of time is virtually impossible to model using this approach. When gas generation is incorporated into the modelling as an overpressure generating mechanism (di Primio and Neumann 2007; Broichhausen et al. 2005) the evolution of overpressure becomes a function of the temperature history of the source rocks. Accordingly, overpressure generation accelerates drastically during gas generation; in the case of Haltenbanken this is in the last few million years when a rapid burial occurred. Figure 6.4.11 compares the two extremes of overpressure evolution calculated assuming exclusively compaction and com-

paction plus gas generation as the overpressure generating mechanisms (see Table 6.4.2). Overpressures are not restricted to shale sequences, but can also occur in evaporite settings, where fluid retention occurs due to the very low permeability of the salt (Urai, Schléder et al. this volume). Examples are carbonate stringers in Palaeozoic salts of the Ara formation in Oman, which vary over short distances between highly overpressured and lithostatic (Fig. 6.4.12). There, retention combined with oil to gas cracking and gas generation probably caused the hard overpressures. Abundant solid bitumen clearly proves the former presence of oil and oil to gas cracking there (Schoenherr et al. 2007a,b).

6.4.7 Effects of glaciation on petroleum systems Gas hydrates have been mentioned above as a possible alternative energy source occurring in shallow sedimentary realms. Their occurrence and especially natural destabilisation and decay have also been implied to exert a prima-

Petroleum systems Figure 6.4.13. Lithostatic and hydrostatic pressure trends and measured pore pressures in carbonate stringers of the infra-Cambrian Ara salt, Oman (after Schoenherr et al. 2007a)

ry control on the atmospheric methane variation observed in ice-core records during the Quaternary (Brook et al. 1999; Chappellaz et al. 1993). Kennet et al. (2003) called this concept the Clathrate Gun Hypothesis, indicating that methane hydrates potentially play a significant role in the rapid warming trends of the last 1 Ma. A rapid or even catastrophic rate of clathrate decay is assumed to be required to release sufficient methane in order to overcome consumption, oxidation and dissolution within the sediments or water column and directly affect the atmosphere (Maslin et al. 2004). Obviously, effects of glacial loading and unloading as well as glacially influenced temperature variations and sea level fluctuations play an important role in changing the stability field of gas hydrates, especially in high latitudes. As an example, in the area of the Mittelplate oil field in Northern Germany (see Magri et al. this volume), gas hydrates were stable during five periods in the past 750,000 years. With depths down to 750 m and thicknesses of up to 500 m these gas hydrate stability fields lasted several 10,000 years (Fig. 6.4.13). The last period of gas hydrate stability ceased just recently with the beginning of the Holocene. However, even if thermally generated gas or microbial gas was periodically trapped in such gas hydrate layers during the Pleistocene, no significant effect of these gas hydrates on the deeper lying Mittelplate oil field can be expected. Little attention has been paid to the effects of glaciations on existing petroleum systems. Here the impact of

ice loading, isostatic rebound and permafrost formation and decay on the underlying petroleum system, especially with respect to enhancing or inhibiting gas migration to the locations of hydrate formation, deserve attention. Cavanagh et al. (2006) indicated that glacial loading and unloading in the Barents Sea, Norway, can have significant effects on the PVT conditions of existing reservoirs and hence systematically suppress gas leakage from reservoirs during glacial periods and enhance leakage during glacial retreat and in interglacial periods. Figure 6.4.14 demonstrates the evolution of reservoir pressure, temperature and depth during the last 1 Ma for the Snøhvit Field in the Barents Sea as determined using basin modelling. During glacial loading pore pressure in the reservoir increases, leading to the compression of the gas phase present in the trap, an increase in the gas density due to enhanced solubilisation of liquid components from the oil leg and, hence, to a decrease in the capillary pressure exerted by the hydrocarbon column on the cap rock. Accordingly, the likelihood of capillary failure of the cap rock decreases. During ice retreat and in interglacial periods pore pressure decreases, the gas expands and becomes drier and the buoyancy force of the hydrocarbon column increases. Based on these results gas leakage, transporting methane to the shallower sites of hydrate formation should be highest during glacial retreat and in interglacial stages. In addition to the effects of ice loading and unloading on ground temperature (Fig. 6.4.14) variable frequencies of low surface temperatures during glaciations can lead to deep spatial temperature distortions in sedimentary ba-

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Figure 6.4.14. Effects of glacial loading and unloading on the pore pressure, temperature and depth evolution of the Snøhvit reservoir taking five glaciation events into account. The uplift of the reservoir formation is due to glacial erosion of the basin

sins. As a result, processes like source rock maturation and hydrocarbon generation rates may vary in time and space. For the example of the oil reservoir in Mittelplate sediments were significantly cooled by up to 7 °C at the beginning of the Elsterian glaciation (Fig. 6.4.15). Even today the reservoir is still about 5 °C colder than expected without glacial cooling. As thermal conductivity of ice is higher than that of fluid water, heat in the sedimentary ba-

sin is more rapidly transported through permafrost. This leads to an additional cooling effect below permafrost layers. For the Mittelplate reservoir this cooling effect reaches up to 1 °C (Fig. 6.4.15). However, a significant influence of glacial conditions on source rock maturation in the Posidonia shale for the example Mittelplate can be ruled out. In consequence, petroleum generation rates were also not lowered perceptibly.

Figure 6.4.15. Effect of cold climate and permafrost on reservoir temperatures in the Mittelplate reservoir

6

Chapter 6.5

Origin and distribution of non-hydrocarbon gases B.M. Krooss · B. Plessen · H.G. Machel · V. Lüders · R. Littke

6.5.1 Introduction The term “natural gas” is commonly used for fuel gases consisting predominantly of methane and minor amounts of higher alkane homologues (C2 – C5 hydrocarbons). Natural gases occurring in sedimentary basins may contain significant amounts of non-hydrocarbon components such as nitrogen, carbon dioxide, and hydrogen sulfide, however. The prediction of non-hydrocarbon contents of natural gases, which may vary from traces to nearly 100% remains one of the major challenges in energy gas exploration. The Central European Basin System (CEBS) hosts several natural gas provinces such as the Southern North Sea region, the Northwest German (with the Lower Saxony Basin and the subsidence area of the Pompeckj Block) and the Northeast German regions, and the Polish basin (Bandlowa 1998). The hydrocarbon gases in the Carboniferous, Permian and Triassic reservoirs of the CEBS mainly originate from coal-bearing Late Carboniferous strata of different sub-basins with temporally and regionally varying subsidence. The relative portions of hydrocarbon and non-hydrocarbon components show regional patterns indicating a

complex interplay of different gas sources, generation and migration processes. While the gas compositions in the British and Dutch part of the CEBS are dominated by hydrocarbons (Cornford 1998; Gras and Clayton 1998), the gases in the German sector of the North Sea and in the eastern part of Germany exhibit high nitrogen contents up to 90% (Müller 1990; Schuhmacher and May 1990; Bandlowa 1998). Regionally, CO2 and H2S constitute major non-hydrocarbon components. CO2 contents vary from <0.1% to >51% in the CEBS, and the highest contents were found in the Triassic Zechstein reservoirs (Bandlowa 1998). Significant concentrations of H2S (up to 35%) occur in the Zechstein reservoirs of the Weser-Ems region (Bandlowa 1998; Mittag-Brendel 2000). The following chapters outline the present state of understanding of the origin and distribution of non-hydrocarbon gases in sedimentary basins.

6.5.2 Nitrogen Molecular nitrogen (N2) regionally constitutes a major component of natural gases in sedimentary basins. High N2 concentrations, sometimes in excess of 90%, and, in consequence, reduced hydrocarbon (HC) contents result in gases with low calorific value (Jenden et al. 1988;

Table 6.5.1. Nitrogen in natural gas accumulations of sedimentary basins world-wide Region

Source rocks

Ural-Volga basin (Prasolov et al. 1990) Californian Great valley Basin (Jenden et al. 1988) Central European Basin (Müller et al. 1976; Hoth et al. 2002)

Vitrinite reflectance of source rock (%)

δ15N (‰)

Reservoir formation

N2 vol (%)

PermoCarboniferous

Permo-Carboniferous

85-100

Tertiary Cretaceous, Late Jurassic metasediments PermoCarboniferous

Tertiary, Cretaceous,

1 to 87

+0.9 to +3.5

1-6

Permian Rotliegend, Triassic, Upper Carboniferous

10 to 95

-5 to +18

Cretaceous

0.1 to 1.2

-19 to -10.7

Pliocene

3 to 33.5

-9 to - 2

West Siberian basin (Zhu et al. 2000)

Cretaceous

0.6

Yinggehai basin (Zhu et al. 2000)

Miocene, Oligocene

0.8-2

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Krooss et al. 1995; Littke et al. 1995; Zhu et al. 2000) and therefore represent a major exploration risk. Nitrogen-rich natural gases occur frequently in the northern and eastern part of the Central European Basin System and therefore this phenomenon has attracted particular attention in industry and academia. High nitrogen contents were found in many other gas reservoirs world wide, such as Estonia, Sweden, Texas, Kansas and Colorado (Krooss et al. 1995 and references cited therein). Natural gases of the Permo-Carboniferous reservoirs of the Ural-Volga basin show nitrogen contents between 85 and 100% (Maksimow et al. 1973). Similar nitrogen contents were found in the north-eastern part of the CEBS (Müller et al. 1976). Overviews of nitrogen-rich natural gas accumulations in different basin were given by Krooss et al. (1995), and Zhu et al. (2000). The ranges of N2-contents and isotopic data in natural gases of selected gas provinces are listed in table 6.5.1. The main hypotheses for the origin of nitrogen-rich natural gas accumulations comprise the following sources or origins: radiogenic nitrogen, atmospheric N2, primordial nitrogen, organic nitrogen, crustal nitrogen. Radiogenic nitrogen can be produced by two nuclear reactions: the exoenergetic reaction of boron or the endoenergetic reaction of carbon with α-particles (helium nuclei). These reactions are, however, considered of very subordinate importance (Krooss et al. 1995). The possibility of a direct contribution of atmospheric nitrogen to nitrogen-rich reservoirs via meteoric water or groundwater has been discussed by various authors (e.g., Marty et al. 1988; Jenden et al. 1988). It is considered unlikely or marginal, particularly for sediments deposited in subaqueous environments. One main argument against significant direct incorporation of atmospheric nitrogen into deep reservoirs is the very high nitrogen/argon ratio in nitrogen-rich natural gas accumulations compared to the atmospheric N2/Ar ratio of 40-80 (Philipp and Reinicke 1982). Jenden et al. (1988) found N2/Ar ratios as high as 22,000 in natural gases from the California Great Valley. The accumulation of primordial nitrogen advocated by some authors relates to the issue of mantle degassing. At the end of the accretion period of the earth about 4.5 billion years ago, the formation of a palaeoatmosphere took place almost completely (99%) within about 50 Ma (Allègre et al. 1987). Since that time the flux of primordial nitrogen decreased continuously as a consequence of cooling of the outer mantle and its contribution to nitrogen-rich natural gas accumulations is considered marginal. Presently, outgassing of primordial nitrogen is probably largely confined to zones of seismic and volcanic activity (Sano et al. 2001) and deep rift systems. Thus, based

on noble gas and isotope analyses, Bräuer et al. (2004) conclude conclude a significant nitrogen flux from the European sub-continental mantle in the Cenozoic Western Eger rift of Central Europe. All other concepts for the formation of nitrogen-rich natural gas accumulations involve the release of nitrogen from organic and/or mineral sedimentary and crustal material by thermal, hydrothermal or chemical processes. This release is envisaged to occur either directly as N2 or as ammonium (NH4+) with subsequent oxidation to N2. Therefore the chemistry of nitrogen in the geo/biosphere is a key factor for the elucidation of the formation of N2-rich natural gas reservoirs.

6.5.2.1 Nitrogen geochemistry and the nitrogen cycle With a total amount of ~4·1021 g the atmosphere contains approximately 50% of the earth’s nitrogen1 (Allègre et al. 2001). The remaining half resides in the lithosphere (crustal nitrogen in sediments ~4 - 6·1020 g) and the mantle. The mean nitrogen concentration of the bulk Earth is ~1.3 ppm. While the estimated average mantle concentration ranges around 1 ppm, crustal rocks are enriched in nitrogen, with a mean concentration in the order of 100 ppm (Wedepohl 1995; Boyd 2001). Due to their abundance, igneous rocks with nitrogen contents of less than 100 ppm constitute the largest reservoir of crustal nitrogen (Blackburn 1983). Sedimentary and metasedimentary rocks may contain 1000 to 2000 ppm of nitrogen and sedimentary organic matter and coals up to 2%. Nitrogen is highly enriched (up to 17%) in the nitrate deposits of the deserts of northern Chile and southern California, which, however, represent very extraordinary geological features. The biosphere plays a key role in the transfer of nitrogen from the atmosphere to the hydrosphere and the lithosphere. Biological fixation represents the main gateway of nitrogen into the biochemical and geochemical nitrogen cycles (Jaffe 1992; Boyd 2001). Only a limited number of organisms, such as blue-green algae, are capable of assimilating atmospheric nitrogen (N2) and thus making it available to the metabolism of other organisms. Because nitrogen is an essential nutrient limiting primary production, it is rapidly recycled in the biosphere. The near-surface biogeochemical nitrogen cycle involves the processes of nitrification (oxidation of NH3 and NH4+) and denitrification (reduction of NO3; Sweeney et al. 1978; Rosswall 1981; Summons 1993; Fogel and Cifuentes 1993; Boyd 2001; Killops and Killops 2005). Amino acids, peptides and nucleic acids are the major nitrogen-containing species in living organisms. These compounds are relatively

1 According to Rosswall (1981), the atmosphere only contains 6 % of all nitrogen on earth. This discrepancy is indicative of the difficulties of quantifying the inventory of the earth’s volatiles.

Origin and distribution of non-hydrocarbon gases

unstable and prone to hydrolysis and recycling. Other, more refractory nitrogen-bearing structures such as pyrrolic or pyridinic units are preserved in the dead organic matter and enter the sedimentary cycle (Fig. 6.5.1a). In the terrestrial system 4% of the nitrogen occurs in living biomass, the remainder forming a large reservoir as dead organic matter (Rosswall 1981). During biological and thermal degradation of organic matter some of the organic nitrogen is liberated as ammonium, which may be incorporated into feldspars and clay minerals (Baxby et al. 1994). Due to their similar ionic radius NH4+ ions substitute potassium ions (K+) at the interlayer sites of illites (Williams et al. 1989; Sucha et al. 1998) (Fig. 6.5.1b). In general, the amount of nitrogen fixed in this way depends on the availability of ammonium during clay mineral formation (Williams et al. 1992) and on the fixing capacity of these minerals (Schroeder and Ingall 1994). Daniels and Altaner (1990, 1993) found that ammonium clays (NH4+–illite, pyrophyllite, rectorite, tosudite and sudotite) are formed from the precursor minerals kaolinite and quartz during the anthracite stage (T >200 °C) of coals. The ammonium originates from the organic material of the coal seams (Ader et al. 1998). According to Daniels et al. (1994), the formation of NH4+bearing illites involves two independent reactions: the release of nitrogen from organic matter and illite formation from kaolinite. Both reactions occur at temperatures >200°C. The formation of ammonium illite can be written as: 3 kaolinite + 2 NH4+ 2 NH4+- illite + 2H+ + 3H2O (6.5.1) During further subsidence and metamorphism ammonium is fixed predominantly in K-micas and feldspars.

in the delta notation relative to atmospheric nitrogen; δ15Natm=0‰) has been used to analyse and compare different sources and identify reaction and transport pathways. Frequently, however, an unambiguous discrimination of different sources is not possible due to an overlap of isotope ranges and complex fractionation processes. Nitrogen in the nitrate deposits of Chile and California has δ15N values ranging from ‑5 to +5‰ (Böhlke et al. 1997). The isotopic composition of nitrogen-fixing algae and bacteria in the euphotic marine environment is similar to atmospheric nitrogen or depleted in 15N (Wada and Hattori 1976; Rau et al. 1987). The 15N enrichment in certain marine sediments probably originates from denitrification processes and is controlled by the availability of nitrate (Smith 1975; Sweeney et al. 1978). The preferential loss of 14NH4+ from minerals due to devolatilisation reactions and/or fluid rock interactions during prograde metamorphism results in a relative enrichment of 15N in the residual nitrogen (Haendel et al. 1986). Thus, the fixed nitrogen isotopic composition in sediments clusters around +1 to +4‰, whereas in high-grade metasedimentary rocks values up to about +10‰ are observed (Haendel et al. 1986; Bebout and Fogel 1992; Boyd and Philippot 1998; Mingram and Bräuer 2001). Magmatic rocks have relatively low nitrogen contents, but S-type granites show an average composition of 45 ppm and δ15N values up to +13‰ (Hall 1999; Boyd et al. 1993). Mantle rocks typically contain small quantities of isotopically light nitrogen. Molecular nitrogen occluded in mid-ocean ridge basalts (MORB) and diamonds commonly has δ15N values in the range of ‑3 to -8‰ (Javoy and Pineau 1991; Cartigny et al. 1997, 1998; Marty and Humbert 1997).

Isotopic composition Nitrogen has two stable isotopes, 14N (99.634%) and 15N (0.366%). The nitrogen isotopic composition (reported

The isotopic composition of nitrogen in natural gases ranges from ‑19 up to +18‰ (Zhu et al. 2000). Zhu et al. (2000) suggest that N2 with very low δ15N values (‑19 to

Figure 6.5.1. a) refractory covalently bonded nitrogen in organic matter; b) NH4+ incorporation into the interlayer sites of clay minerals (after Williams et al. 1989)

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Figure 6.5.2. Nitrogen isotopic composition of different nitrogen reservoirs of the earth. Modified after Snyder et al. (2003): (1) Zhu et al. 2000 (2) Stiehl and Lehmann 1980; Sweeney and Kaplan 1980; Rau et al. 1987; (3) Mingram and Bräuer 2001; (4) Boyd et al. 1993; Hall 1999; (5) Javoy and Pineau 1991; Cartigny et al. 1997; Bräuer et al. 2004 (6); Marty and Humbert 1997; (7) Jenden et al. 1988; Sano et al. 2001; Fischer et al. 2002; (8) Böhlke et al. 1997

‑10‰) originates from immature organic matter (vitrinite reflectance VRr<0.6%), whereas δ15N values between ‑10 and ‑2‰ reflect an origin from mature organic matter (VRr 0.6 to 2.0%) and very heavy δ15N values (+4 to +18‰) indicate post-mature organic source material. However, the nitrogen isotopic composition of natural gases is not only influenced by the sources but is also affected by the release mechanisms and migration processes. Non-equilibrium or kinetic fractionation processes associated with diffusion, adsorption and phase transfer (solution/exsolution) have been postulated repeatedly to influence the isotopic composition of N2 and cause a relative depletion of 15N at the migration front (Müller et al. 1976). Zhang and Krooss (2001), after analysing different natural gas migration scenarios conclude, however, that these isotope fractionation effects are usually overridden and wiped out by volume transport (bulk flow) processes.

6.5.2.2 Nitrogen in the NGB Nitrogen in natural gases Natural gases from Carboniferous, Permian and Triassic reservoirs of the North German Basin are mainly sourced from coal-bearing strata and marine shales of Carboniferous age (Gras and Clayton 1998). The gas compositions range from predominantly hydrocarbons with only a few percent nitrogen in the western part up to 90% nitrogen in

the eastern part of the basin (Fig. 6.5.3) (Bandlowa 1998; Gerling et al. 1997). The isotopic composition of N2 in the nitrogen-rich gas reservoirs ranges between +3 and +10‰ while the nitrogen-lean reservoirs exhibit a much greater variability in N-isotopic composition from -5 to +18‰ (Müller et al. 1976; Gerling et al. 1997). The broad range in the isotopic composition of the latter reservoirs is attributed to different sources and migration paths (Hoth et al. 2002). In the eastern part of the North German Basin (northern Germany and Poland), Westphalian coals are essentially absent and the deeply buried Permian strata are underlain by Namurian shales with relatively low contents of organic matter. The absence of the coal layers with their typically high methane generation potential could explain the abundance of Permian Rotliegend gas reservoirs with high nitrogen contents (>90% N2) in the deepest part of this basin (Gras and Clayton 1998). These nitrogen-rich gases have also relatively high helium contents with values up to 0.2 % (Bandlowa 1998). The N2/He ratios range around 300 and are thus significantly higher than mantle N2/He ratios of mid-ocean ridge basalt (MORB) glasses (20 to 200; cf. Jenden et al. 1988). The 3He/4He ratios2 of gases from deep fluids in the Rotliegend of the eastern part of the NGB range around 2·10-8 (~0.02 Ra) providing a clear indication of crustal origin (Wiersberg et al. 2004). Because influx of nitrogen from the mantle would also be expected to carry significant amounts of mantle helium, the molecular nitrogen in this area is very likely to derive from crustal sources.

2 The atmospheric 3He/4He ratio, Ra, is 1.38·10-6; mantle gases, and gases in mantle-derived basalts have higher ratios (5 to 20 Ra). Crustal rocks have 3He/4He ratios of 0.009 to 0.056 Ra due to 4He generated by radioactive decay of uranium and thorium.

Figure 6.5.3. Distribution of Late Carboniferous sediments and nitrogen contents in Rotliegend gas reservoirs of the Central European Basin System; Hoth et al. (2002); Corcoran and Clayton (2001), TNO, BGS, GEUS, BGR, PIG – European Gasatlas, Lokhorst et al. (1998)

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Figure 6.5.4. (a) Organic carbon vs. total nitrogen contents (wt. %) of Palaeozoic black shales, coals and anthracites, fluvial and deltaic siltstones and one Tertiary lacustrineblack shale from the CEBS (after Littke et al. 1995); (b) Atomic nitrogen/carbon ratio vs. maximum vitrinite reflectance of coals (after Boudou et al. 2008)

Nitrogen in Palaeozoic rocks of the CEBS While Namurian and Westphalian coals of the CEBS have nitrogen contents of up to 3.5% (Fig. 6.5.4), the nitrogen contents of carbonaceous shales reach 2800 ppm with an inorganic fixed nitrogen portion in the form of ammonium up to 90% (Mingram et al. 2005). The highest nitrogen values were found in Namurian pelagic shales whereas the limnic to paralic Westphalian shales show lower nitrogen contents (Hoth et al. 2002). The stable isotopic composition of fixed nitrogen in all studied Carboniferous sediments of the CEBS extends over a narrow range between +1‰ and +3.5‰ (Mingram et al. 2005). These variations were found in shales of varying maturity (VRr from 1.5 to 6%) and appear to reflect differences in the sedimentary facies and depositional conditions rather than effects of thermal history. Between 60 and 90% of this nitrogen are fixed in clay minerals or illites as NH4+. Cambrian to Carboniferous sediments of the NGB and surrounding areas from several cores and outcrops show ammonium nitrogen contents between 400 and 1500 ppm (Mingram and Bräuer 2001; Mingram et al. 2005). The range of δ15N isotopic composition is relatively narrow with values between +1 to +4‰. The lowest fixed nitrogen contents with the highest δ15N values were found in PermoCarboniferous magmatic rocks (<100 ppm, +6 to +15‰). Permian shales represent a mixture of magmatic, sedimentary and metamorphic rock detritus with intermediate fixed nitrogen contents (100 to 500 ppm, δ15N +2 to +6‰). The sequence of Cambrian to Carboniferous sediments in the NGB comprises a depth interval of more than 5 km.

The release of only 300 ppm of nitrogen from a rock volume of 500 km3 would generate 320 billion m3 of N2. This quantity is in the order of the gas volume of the large Salzwedel-Peckensen reservoir of the Altmark province in central Germany.

6.5.2.3 Mechanisms and conditions of nitrogen release Several studies have investigated the potential processes resulting in the formation of the nitrogen-rich gases in the CEBS (e.g., Müller et al. 1973, 1976; Müller 1990; Scholten 1991; Krooss et al. 1995, 2005; Littke et al. 1995; Hoth et al. 2002; Mingram et al. 2003, 2005). Some of these studies initially focused on organic nitrogen, especially in coals, as a probable source of nitrogen in the CEBS (Krooss et al. 1993, 1995). As stated above, nitrogen in clastic sediments can occur both in organic and inorganic compounds. Depending on the depositional environment and preservation conditions these compounds will be present in different proportions. During subsidence and burial of the sedimentary layers the fixed nitrogen is successively liberated as a result of increasing thermal stress and changing geochemical conditions (cf. Cooper and Evans 1983; Whelan et al. 1988; Everlien and Hoffmann 1991; Jenden et al. 1988; Krooss et al. 1995; Zhu et al. 2000; Mingram et al. 2005). The phases of nitrogen release from sedimentary rocks extend from early diagenesis to late metamorphism and vary in intensity with time, depending on the stability of the nitrogen-bearing organic and inorganic components. Although it is widely recognised that the release or devolatilisation of

Origin and distribution of non-hydrocarbon gases Figure 6.5.5. Pyrolytic release of N2 from Carboniferous coals of different rank. Percentage values indicate vitrinite reflectance

nitrogen from sedimentary sequences plays a predominant role in the lithospheric nitrogen cycle, the conditions under which these processes occur are still poorly understood. Temperature and fluid-rock interactions are, however, generally considered to be the main controlling factors. Various approaches have been used during recent years to elucidate the mechanisms and conditions of nitrogen release in sedimentary basins and arrive at an improved understanding of these processes as related to the formation of nitrogen-rich natural gas reservoirs. These comprise: (i) the assessment of the thermal stability and decomposition kinetics of nitrogen-containing species in sedimentary rocks and coals by pyrolytic methods (ii) the detection of nitrogen release processes and isotopic composition in sedimentary and metamorphic rocks, and (iii) the analysis of fluid inclusions. Release of nitrogen from coals Pyrolytic methods are commonly used to obtain information on the thermal stability of solids, the amounts and composition of the products and the kinetics of the decomposition reactions. Both, open and closed system pyrolysis has been used extensively to elucidate the mechanisms and kinetics of petroleum and natural gas generation from sedimentary organic matter. Kinetic parameters derived from laboratory pyrolysis are routinely used in petroleum system analysis to reconstruct the timing and intensity of oil and gas generation. First attempts to apply these concepts to the generation of molecular nitrogen were made by Klein and Jüntgen (1972) and Jüntgen and Klein (1975) who conducted non-isothermal pyrolysis experiments to study the thermal release of N2 and methane from Carboniferous coals.

Coals and sedimentary organic matter (SOM) are the main sources of hydrocarbon gases in the CEBS. They also exhibit the highest concentrations of fixed nitrogen in sedimentary sequences and have therefore also been considered as potential sources for N2-rich natural gas accumulations. Figure 6.5.4a shows a general correlationbetween organic carbon (Corg) and total nitrogen contents (Ntot) of black shales, coals and siltstones, which is indicative of the role of organic matter in the sedimentary nitrogen cycle. The considerable scatter in the Ntot vs. Corg plot of coals is due to the variability of the coal composition and the rank-dependence of the nitrogen contents. During the process of thermal maturation, sedimentary organic matter successively loses nitrogen, as indicated by a decrease in the atomic N/C ratio with increasing vitrinite reflectance (Fig. 6.5.4b). Due to the inherent variability of the nitrogen content of lower rank organic matter (VRr <1.5%), this trend is clearly discernible only in the semi-anthracite to meta-anthracite range (VRr =2.0 – 6.5%). Boudou et al. (2008) investigated the speciation and fate of organic nitrogen before and during metamorphism. They conclude that with increasing metamorphic grade, pyridinic and pyrrolic nitrogen structures gradually disappear and the residual nitrogen is fixed in condensed, partially aromatic systems where N is covalently bonded to three C atoms. Isotope analyses of the fixed nitrogen suggest that during this late-stage maturation Norg is eliminated without significant nitrogen isotope fractionation. The methane generation potential of coals is at least ten times higher than the N2 generation potential. Therefore higher N2/CH4 ratios of coal-derived gas can be accounted for only if (i) the generation of methane and N2 occurs sequentially and (ii) the reservoir contents represent only

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Figure 6.5.6. Thermal release patterns of N2 from different Carboniferous shales

the composition of the latest gas generated (dynamic reservoir concept; cf. Krooss et al. 1995). The pyrolysis experiments of Klein and Jüntgen (1975) and Stiehl and Lehmann (1980) indeed revealed that N2 was liberated from coals at significantly higher temperatures than methane. Krooss et al. (1993, 1995) performed similar nonisothermal pyrolysis experiments on a maturity sequence of Westphalian coals (VRr = 0.78 – 6.1%). The pyrograms of this coal sequence (Fig. 6.5.5) show a maximum in N2 release rate (generation peak) in the 700 – 800 °C temperature interval. With increasing coal rank this maximum shifts towards higher temperatures while its intensity decreases. For the meta-anthracite (VRr=6.1%) N2 release only occurs at temperatures above 1000 °C. These observations are in line with changes in N/C ratio with rank (Littke et al. 1995, Boudou et al. 2008) showing a successive loss of nitrogen from the coal during thermal maturation. In terms of speciation, the thermally “labile” nitrogen species decomposing in the 600 – 900 °C temperature range are likely to represent the pyrrolic and pyridinic structures, while the more stable nitrogen components decomposing only at temperatures well above 1000 °C would correspond to the condensed, partially aromatic structures proposed by Boudou et al. (2008). To investigate to what extent thermal gas generation from coals could result in nitrogen-rich natural gases, reaction-kinetic parameters for the generation of metha-

ne and N2 were determined for a Westphalian coal of the Ruhr area and subsequently used in a numerical basin analysis of the North German Basin. In this study, quantitative aspects of the formation of natural gas reservoirs such as reservoir size, source rock, drainage area, timing of deposition of seals and trapping were examined. It became clear that large amounts of nitrogen have been generated from organic matter in the CEBS, especially in areas of high thermal maturity of the coal-bearing sequence. However, due to the overlap of the periods of methane and N2 generation, none of the conceivable gas generation scenarios reproduced the high N2 concentrations observed in some of the gas reservoirs of the study area. The predicted N2 contents did not exceed 50% even when selective (non-cumulative) trapping and accumulation of only the latest, N2-rich, gas charge was considered. It was therefore concluded that coal-bearing strata alone cannot account for the high N2 contents of the gas reservoirs in the NGB. Release of inorganic fixed nitrogen Numerous studies have documented the occurrence of nitrogen fixed as ammonium in illites and feldspars of (carbonaceous) sedimentary and metamorphic rocks (Higashi 1978; Cooper and Evans 1983; Juster et al. 1987; Williams et al. 1989, 1995; Boyd 1997; Sucha et al. 1994). The release of ammonia from these silicates may occur by

Origin and distribution of non-hydrocarbon gases

Figure 6.5.7. N2 pyrogram obtained from non-isothermal open-system pyrolysis (heating rate 0.5 °C min) of a Namurian A shale from NW Germany and isotopic composition of N2 determined by GC/IRMS

thermal decomposition, cation exchange reactions and/or oxidation. According to Eugster and Munoz (1966) the thermal decomposition of ammonium muscovite occurs at temperatures above 400 °C whereas cation exchange and oxidation proceed at lower temperatures. Thermal release of inorganic nitrogen Different studies performed on metamorphic rocks showed that during the thermal devolatilisation nitrogen will be released as N2 or NH3 with a preferred release of 14 N and a relative enrichment of 15N in the remaining nitrogen (Haendel et al. 1986; Bebout and Fogel 1992; Mingram and Bräuer 2001). Everlien (1990) monitored the release of ammonia and N2 from nitrogen-bearing clay minerals by non-isothermal pyrolysis experiments (2.5 °C/min up to >1000 °C). These experiments were performed with various artificial ammonium clays. Ammonium montmorillonites and illites were found to produce predominantly ammonia and only little or no N2, while N2 was the main pyrolysis product from ammonium vermiculite. Oh et al. (1993) investigated the thermal decomposition of buddingtonite, an ammoniumbearing feldspar, and found that ammonia evolved at temperatures between 400 and 900 °C. H2 and N2 were lib-

erated at temperatures above 500 °C as a consequence of NH4+ decomposition. Applying reaction kinetic parameters derived from laboratory experiments to burial and thermal histories of sedimentary sequences in northern Germany, Everlien and Hoffmann (1991) concluded that nitrogen release from smectites and illites should occur at a depth of about 5000 m. Gas release from ammonium vermiculites is expected only at burial depths of 10,000 m. Using the same open-system pyrolytic method as for coals, Krooss et al. (2005) analysed the thermal release of molecular nitrogen and other volatile components (methane, carbon dioxide and carbon monoxide) from Palaeozoic carbonaceous shales and siltstones with varying TOC contents. These experiments revealed detailed insight into the presence of nitrogen-bearing compounds and their thermal stability. The N2 pyrograms show reproducible release patterns that are characteristic for certain facies types and depositional environments (Fig. 6.5.6). Thus, the marine Namurian A-B (cnA-B) shales exhibit an intense release of N2 at relatively low temperatures. Comparison of the pyrograms of the bulk shale samples and the corresponding kerogen, isolated by HF treatment, revealed that the major portion (60%) of the N2 originated from the silicate phase while only smaller amounts of N2

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liberated at high temperatures (>900 °C) could be attributed to organic matter (Krooss et al. 2005). The N2 release observed during these laboratory experiments can be formally described by reaction kinetic parameters assuming sets of parallel first-order reactions. The applicability of these kinetic data to geologic burial and heating histories is still debated. In order to investigate the isotope geochemistry of nitrogen fixed in sedimentary rocks, the open-system pyrolysis technique was combined with on-line 15N isotope analysis of the liberated nitrogen. The N2 pyrogram of a Namurian A shale with the pertaining isotopic composition is shown in figure 6.5.7. During the main phase of N2 generation, the δ15N values increase but then exhibit a continuous trend towards lighter values beyond 700 °C. Meanwhile, various different isotope trends have been observed for different shale samples indicating mixtures of precursor species with different thermal stabilities and isotopic compositions. These pyrolytic studies revealed that the marine Namurian A shales of the NGB are enriched in a thermally labile nitrogen-containing silicate, which may liberate considerable amounts of molecular nitrogen and/or ammonium at great burial depth and correspondingly high temperatures. The role of fluid-rock interactions and deep fluid transport processes in this release process still remains to be clarified. Release of ammonium (NH4+) from minerals by fluid/ rock interactions The bulk of the nitrogen (up to 90%) in the pre-Westphalian Palaeozoic sediments of the NGB is fixed in the

form of ammonium in clay minerals of thick shale layers. Therefore, the fate of this species during basin evolution plays a major role in the process of the formation of nitrogen-rich gas accumulations. While the thermal release of N2 from Palaeozoic shales was clearly established to start at temperatures around 450 °C under laboratory conditions, the question arises whether temperatures of max. 300 °C reached during subsidence are sufficient to achieve this release on the geologic time-scale. This issue relates to the kinetics of the decomposition reactions and to the thermodynamic stability of the as yet poorly characterised nitrogen-containing compounds. Another major unknown is the influence of formation fluids on the release process. The behaviour of ammonium during substitution for + potassium in clay minerals, the release as NH3/NH4 , the + migration as NH4 with fluids and oxidation to N2 was investigated by mineralogical and geochemical studies on potential source rocks, clay minerals, and fluids (Mingram et al. 2005; Lüders et al. 2005). Systematic analysis of total nitrogen and NH4+ nitrogen potential and the nitrogen isotopic composition trends in NE Germany resulted in the identification of a location (Parchim) where the ammonium contents of Palaeozoic shales were much lower (~460 ppm) than in surrounding areas (>1000 ppm). The ammonium nitrogen in these shales was enriched in 15N by up to +5.6‰ with respect to other Palaeozoic shales from nearby wells. These observations suggest a significant release of nitrogen at this location (Fig. 6.5.8; Mingram et al. 2005). Calculation of nitrogen loss and isotopic fractionation indicates that more than 30% of the nitrogen was released as ammonium by thermal impact during burial or

Origin and distribution of non-hydrocarbon gases

by fluid/rock interactions with highly saline brines. This nitrogen loss associated with a shift in δ15N for the residual nitrogen suggests a NH4+-NH3 exchange reaction in the temperature range of 250-350°C. Although the shales from the entire study area had similar thermal maturities (VRr values ranging from 3.4 to 5.1%), samples from the Parchim well are affected by fluid/rock interactions and show indications of carbonate displacement and chloritisation of illites. Significant evidence was found for a release of NH3/ NH4+ from the illites by cation exchange processes with K-bearing brines (Mingram et al. 2005). Preliminary results of K-Ar dating on NH4+-bearing illites indicate that illite did not form only during burial diagenesis but also during hydrothermal alteration that caused the NH4+ or K+ exchange processes. The migration mechanism of NH4/NH3 liberated from the shales differs substantially from the migration of N2 (either dissolved or as a free gas phase). Ammonia is highly soluble and can migrate rapidly through the water-saturated pore space. During migration NH4+ may be reincorporated in authigenic illites or oxidised to N2. Secondary authigenic illites in Permo-Carboniferous felsic volcanic rocks (primary NH4+-N poor) indicate a hydrothermal overprint by external NH4+-enriched fluid during the Triassic/Jurassic transition. Ammonium may be oxidised to N2 when the migrating fluids pass through hematite-rich Rotliegend strata that occur abundantly in the North German Basin. The large amounts of hematite in Rotliegend sediments and volcanic rocks represent a high oxidation potential for the oxidation of NH4+ to N2. According to Getz (1980) the reaction can be formulated as follows: 9Fe2O3 + 2NH3 → 6Fe3O4 + N2 + 3H2O

The release of N2 from these different types of fixed nitrogen is controlled by their thermal stability and the subsidence and thermal evolution of the corresponding sedimentary sequences. Pyrolytic studies have revealed the wide range of thermal stabilities of organic and inorganic nitrogen-bearing species in Palaeozoic rocks and their relation to specific depositional environments. The release of N2 from coals occurs at high temperatures and is associated with the formation of considerable amounts of methane. Thus, it cannot readily account for high (>>50%) N2 contents observed in many natural gases of NE Germany. Release of nitrogen from mineral species occurs at lower temperatures and is most likely also influenced by fluid-mineral interactions. While, based on laboratory experiments and regional geochemical and isotopic studies, a conceptual framework for the release of N2 and NH4+ from the crustal inventory of fixed nitrogen could be established, the time/ temperature conditions under which these processes occur still remain poorly constrained. Integration of these geochemical concepts into the 3D basin analysis of different gas provinces is one promising approach to close this gap. Furthermore, earlier observations related to the gas and isotope geochemistry of N2-rich gases should be reconsidered in the light of the new experience gathered. For instance, correlations of N2 and helium contents reported by various authors could be related to simultaneous release of both gases from deep crustal sources. Further studies should investigate how these features are related to thermal and subsidence history as well as thickness and composition of deep sedimentary layers.

(6.5.2)

The isotopic composition of the molecular nitrogen will be controlled by the isotopic composition of ammonium in the formation fluids. The strongest isotopic fractionation is caused by evaporation processes and changes in pH, T, and p. These result in a partial release of isotopically light ammonia and enrichment of the residual ammonium in 15 N. This process could provide a further explanation for differences in the nitrogen isotopic compositions of the NE German gas reservoirs.

6.5.2.4 Summary Various pieces of evidence indicate that the N2-rich gases of the CEBS and other sedimentary basins originate predominantly from crustal sources. Sedimentary organic matter and coal have the highest fixed nitrogen concentration. Most of the crustal nitrogen resides, however, in inorganic compounds, most likely ammonium clays and feldspars, of thick shale sequences.

6.5.3 Carbon dioxide Free carbon dioxide (CO2) gas is found in natural gas reservoirs worldwide in varying percentages ranging from traces up to 100%. In higher concentrations CO2 is an unwelcome diluent of hydrocarbon gases and in combination with water acts as a corrosive agent. Oil and gas exploration has a strong interest in avoiding CO2, or at least knowing when to expect it in order to assess and minimise the economic risk. High-purity natural CO2, on the other hand, is of economic value and produced in significant quantities for food/beverage and cryogenic use. Natural gas reservoirs with high CO2 contents are of interest as analogues for the geological storage (sequestration) of CO2 in the context of greenhouse emission reduction (Pearce et al. 2004). CO2 occuring in sedimentary basins can originate from inorganic and organic sources (Table 6.5.2). Very high CO2 concentrations in natural gas reservoirs are usually related to volcanism and magmatism while CO2 admix-

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gases (e.g., 3He/4He) to assess the origin of CO2-rich natural gases.

6.5.3.1 CO2–rich natural gases in the CEBS

Figure 6.5.9. Isotopic composition of CO2 from different sources (after Dai et al. 1996)

tures to natural gases up to 10-15% can be due to the decomposition of sedimentary organic matter (Dai et al. 1996; Imbus et al. 1998; Wycherley et al. 1999). Based on a comprehensive survey of natural gases in China, Dai et al. (1996) proposed a classification scheme for CO2-containing natural gases taking into account the chemical and isotopic composition. As shown in figure 6.5.9, high CO2 concentrations and δ13C values higher than -8 to -10‰ are clear indications of inorganic sources whereas decomposition of organic matter results in CO2 concentrations not higher than ~15% but δ13C values lighter than ‑10‰. In recent years various studies have attempted to use the abundance and isotopic composition of associated noble

Permian Zechstein carbonate reservoirs (Stassfurt-Carbonate, Ca2) in the Lower Saxony Basin (LSB) of Northern Germany contain locally high (>90%) concentrations of CO2, constituting a substantial risk for hydrocarbon exploration (Petmecky 1998). Up to now no unambiguous relation between the amount of CO2 and reservoir geological or geochemical parameters has been established and, in consequence, reliable predictions on prospective drill locations are difficult to make. The “CO2-province” is located approximately 60 km north-west of the city of Hannover (Fig. 6.5.10) in the central part of the LSB, which represents an intra-plate crustal segment in NW Germany some 300 km long and about 60 km wide. It is an ENE-WSW striking basin crosscutting the older, northernmost fringe of the WSW-ENE striking Variscan orogenic front (Brink et al. 1992; Maystrenko et al. this volume; Voigt et al. this volume). The LSB subsided between the Pompeckj Block in the north and the Rhenish Massif in the south in Late Jurassic to Early Cretaceous times (Boigk 1981). Inversion phases during Late Cretaceous and Early Tertiary led to an uplift of the sediments (Betz et al. 1987; Senglaub et al. 2005; Littke et al. this volume). The pre-salt basement is intensely structured by normal and reverse faults into a pattern of horst and graben structures as a result of extensional tectonics (Rockenbauch et al. 1998). The horsts host Northern Germany’s sour gas reservoirs (Kockel et al. 1994), which are associated with the “Zechstein 2 Carbonate” or “Ca2”. These carbonates are part of the “Basal Zechstein” consisting of the Werra-anhydrites (A1) underlying the Ca2 and the Stassfurt-anhydrites

Table 6.5.2. Organic and inorganic sources of CO2 in sedimentary basins (after Imbus et al. 1998) Inorganic CO2 sources mantle (vs crustal) emanations

References Giggenbach et al. 1993; Jenden et al. 1993; Poreda 1995; Wycherley et al. 1999

carbonate metamorphism and hydrolysis

Grabowski et al. 1985; Clayton et al. 1990; Baric et al. 1991

chemical equilibria among feldspar, clay and carbonate minerals in siliciclastic and carbonate reservoirs

Smith and Ehrenberg 1989; Reaves and Sulaeman 1994

Alteration of organic material diagenesis and maturation of kerogen, particularly coals

Andresen et al. 1994; Killops et al. 1994; Seewald et al. 1994

decarboxylation of acids formed by hydrous pyrolysis

Kharaka et al. 1993

thermochemical sulfate reduction

Sassen and Moore 1988; Machel et al. 1995b

microbial activity

Martini et al. 1997

Origin and distribution of non-hydrocarbon gases

Figure 6.5.10. (a) The Lower Saxony Basin (LSB) located in the southern part of the Zechstein Basin, NW Germany (modified after Füchtbauer and Peryt 1980). (b) CO2-gas concentration map of the Lower Saxony Basin (modified after Lokhorst et al. 1998). (c) Ca2 sour gas field “Wiehengebirgsvorland” with sediment and gas sampling locations; LSW=“lowstand wedge”. (d) Stratigraphy of the Basal Zechstein (modified after Strohmenger et al. 1996). The Stassfurt Carbonate (Ca2) layer is marked in grey

(A2) overlying the Ca2 reservoir and acting as a seal (Fig. 6.5.10d; Strohmenger et al. 1996). The anhydrite-carbonate-anhydrite triplet forms the uppermost part of the presalt basement in NW Germany. Interestingly, the overlying Triassic (Buntsandstein) reservoirs generally do not contain high CO2-concentrations (Petmecky 1998). The area underwent several phases of extensional and compressional tectonics. As a consequence, the basal Zechstein, as a part of the pre-salt basement, was intensely faulted and fractured.

The natural gas composition of the Ca2 reservoirs of the central part of the LSB varies strongly even within individual structural blocks and over short distances (Gerling et al. 1999). In the Ca2 carbonate, H2S is another important non-hydrocarbon component besides CO2 and N2. Three out of twelve gas-bearing wells in this area show high CO2 concentrations from 74% up to 91%. The complex gas situation is the result of the structural and thermal evolution of the LSB. Central parts of this basin

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ary/Quaternary times, whereas the methane in the reservoirs is generally older. CO2 migration seems to be related to the presence of deep-reaching faults, which are known to be present in the area.

Figure 6.5.11. CO2 concentration data of reservoir gases sampled in wells of the LSB. Data represent possible mixing situations of fluids from different origins. End-member “A”: low CO2 concentration of biogenic origin (δ13C~-20‰). End-member “B”: high CO2 concentration from carbonate decomposition (δ13C~0‰). Theoretical mixing curves represent different ratios of CO2 concentrations of two mixing fluids with δ13C values for end-member gases “A” and “B”, respectively (after Fischer et al. 2006)

underwent intense maturation due to deep burial followed by inversion and erosion during the Late Cretaceous (Petmecky et al. 1999; Littke et al. this volume). Thus, the major phase of methane generation and accumulation from Carboniferous and other Palaeozoic source rocks was the Early Cretaceous. Methane accumulations only exist where they survived the later uplift and erosion phase. Whereas methane generation and accumulation ceased after the onset of inversion, carbon dioxide accumulation has to be regarded as a much younger phenomenon. Carbon dioxide-enriched waters are well known from many present-day springs in the area and have been investigated with respect to their carbon isotope values. Most springs contain carbon dioxide with δ13C values between ‑1 and ‑4‰, although more negative values (up to ‑21‰) are also occasionally reported (Petmecky 1998 and references therein). These values can be compared to those measured on large gas fields in the Zechstein carbonates, which typically range from 0 to +5‰, with some exceptionally low values in the northern part of the CO2-province (down to ‑16.5‰; Petmecky 1998). The overall similarity with respect to isotope values of carbon dioxide between recent springs and deep reservoirs suggests that the CO2 in the reservoirs is also related to young or even ongoing processes, i.e., that filling occurred mainly in Terti-

However, the generation and migration processes resulting in the selective accumulation of individual nonhydrocarbon components are not yet fully understood. The issue is further complicated by the possibility of in-situ generation of H2S and CO2 by thermochemical sulfate reduction (TSR) within the Zechstein anhydrites and carbonates (Steinhoff 1998; Steinhoff and Strohmenger 1996). Fischer et al. (2006) conducted a detailed study on reservoir diagenesis, mineralogy and isotopic composition of the Ca2 reservoirs of the LSB. They investigated the isotope geochemical relationship between authigenic carbonate minerals and related variations of CO2 gas contents in reservoir rocks in order to understand the processes of CO2 migration and accumulation in the CO2 province of the NW German Basin. Carbonate sediments that are affected by dissolution, recrystallisation and/or cementation during rock formation record their diagenetic history by the mineral and geochemical composition of the carbonate components. The study revealed significant differences in mineralogical, geochemical, and in particular stable isotope data (13C, 18O) of fracture- and cleat-filling minerals within the Stassfurt carbonate for methane-rich reservoirs and carbon dioxiderich reservoirs.

6.5.3.2 Origins of CO2 in Ca2 reservoir rocks CO2 contents in the Ca2 reservoirs of the LSB “Wiehengebirgsvorland” area range between 0 and 90% (Lokhorst et al. 1998). Carbon isotope data of authigenic carbonate minerals from Ca2 reservoir rocks suggest that a 12C-enriched carbon source has prevailed in the Ca2 pore water during diagenesis. As the fractured carbonates are most probably related to tectonic events (Strohmenger et al. 1996), local pressure release followed by CO2 degassing processes may have triggered carbonate precipitation within open fractures. Pressure release was probably most pronounced during the Cretaceous phase of deep burial and also during the Tertiary. Pore pressure calculations indicate that high overpressures developed during rapid Cretaceous burial. Fluid retention below and within the evaporitic Zechstein units will have had an effect and may have caused maintenance of overpressures for long periods of time. The carbon isotope composition of CO2 within the Ca2 reservoir rocks appears to be related to the CO2 content of the reservoir gas. The most positive δ13C values were

Origin and distribution of non-hydrocarbon gases

6.5.4 Hydrogen sulfide (H2S)

measured for reservoir gases with high CO2 contents whereas low CO2 concentrations are characterised by negative carbon isotope values. This suggests a mixing process between two carbon source end-members, which are represented by fields A and B, respectively in figure 6.5.11. The two fields also correspond roughly to the carbon source end-members assumed for diagenetic carbonate formation. The isotopically light CO2 (δ13C~-20‰; “A”) most probably derived from the degradation of organic matter. The second “heavy” CO2 source (δ13C ~0‰; “B”) was locally admixed to or replaced the formerly prevailing 13C-depleted CO2. Variable mixing ratios between these two CO2 end-members are indicated in figure 6.5.11. These mixing ratios probably determine the relationship between concentration and isotope compositions of the present-day CO2 in the Ca2 reservoirs.

Hydrogen sulfide (H2S) occurs as a non-hydrocarbon component of natural gases most commonly in geological settings of carbonates and evaporites, but also in siliciclastics. Most hydrogen sulfide forms via redox-reactions between sulfate and hydrocarbons (and many other organic compounds). Such reactions can take place with the mediation of microbes (mainly bacteria) or inorganically (abiologically). Microbially/bacterially mediated sulfate/hydrocarbon redox-reactions are commonly called BSR (bacterial sulfate reduction), while abiological sulfate reduction is called TSR (thermochemical sulfate reduction).

The absolute quantities of CO2 derived from these two sources are not known. Based on the gas samples analysed, 80% of the total reservoir CO2 could derive from a source with a positive carbon isotope values around 0‰. It is well known that CO2 reduction during (early) microbial methanogenesis results in an enrichment of the residual CO2 in 13C (Claypool and Kaplan 1974; Irwin et al. 1977; Whiticar et al. 1986). Isotopically heavy CO2 is also produced from carbonate decomposition processes (Clayton et al. 1990). For the Ca2 rocks, carbonate decomposition processes are much more likely (Smith and Ehrenberg 1989; Cathles and Schoell 2007) because microbial CH4 was not found in the reservoir gases and biogeochemical processes are unlikely to occur at the high temperatures prevailing during late diagenesis. There is evidence, however, of significant pressure solution by numerous stylolites observed in several drill cores. This pressure-induced carbonate dissolution is a likely source for 13C-enriched CO2 admixed to the early organic-derived CO2 within the Ca2 reservoir rocks of the CO2-province in the NW-German Basin.

BSR and TSR are important for at least two reasons. Firstly, they take place in a great variety of sedimentary and diagenetic environments, ranging from waterlogged soils to marine sediments down to a water depth of about 10 km, from sedimentary base metals deposits to reefal carbonates and evaporites that have been subjected to diagenetic processes during burial (see summary in Machel 1989). Reaction products of BSR and TSR have been found worldwide in sediments and rocks ranging in age from the late Proterozoic to the Recent. Secondly, BSR and TSR have significant economic importance. In particular, many deeply buried carbonate rocks contain the common reaction products hydrogen sulfide and/or elemental sulfur), which are valuable resources for sulfuric acid and elemental sulfur production. Also, in some locations, particularly salt dome cap rocks and reefs, reaction byproduct transition and base metal sulfides have been found in economically significant concentrations (e.g., Price and Kyle 1983), Mississippi-Valley-Type lead-zinc sulfide deposits (e.g., Powell and Macqueen 1984), and even some emerald deposits (e.g., Giuliani et al. 2000).

6.5.3.3 Summary In the CO2 province of the LSB an organic source producing an end-member CO2 ( δ13C ~ ‑20‰) is present in relatively low concentrations. High CO2 concentrations derive from an end-member source with a δ13C value near 0‰, most likely Zechstein carbonate, which, upon decomposition, produces isotopically heavy CO2 (Fischer et al. 2006). The Ca2 carbonates were subjected to multiple diagenetic events that occurred during different post-depositional periods. The present reservoir quality is controlled by five major diagenetic events: (i) meteoric leaching, (ii) dolomitisation, (iii) displacive and replacive anhydritisation, (iv) dedolomitisation (calcitisation) and (v) halite cementation (Strohmenger et al. 1996, 1998).

6.5.4.1 Overview

H2S concentrations vary widely, in some provinces such as western Canada between about 1 and 90%, and are one of the most difficult geochemical parameters to predict. For example, variable and in parts rather high concentrations of hydrogen sulfide are found in natural gases of Permian and Carboniferous sequences of Western Canada and Kazakhstan (Karachaganak and Tengiz fields), and from Early Triassic reservoirs from the northeastern Sichuan Basin in SW China (Table 6.5.3).

6.5.4.2 Microbial and thermochemical sulfate reduction The association of dissolved sulfate and petroleum is thermodynamically unstable in diagenetic environments

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Table 6.5.3. Carbon dioxide and hydrogen sulfide contents of gases in the sour gas provinces of the North German Basin, western Canada and SW China (after Fabian 1963; Machel et al. 1995b; Li et al. 2005; Zhu et al. 2005) Well North German Basin (Permian Zechstein gases) Düste Z1 Düste Z3 Buchhorst Z2 Barenburg Z1 Alberta, western Canada (Devonian and Mississipian) Pincher Creek Jumping Pound Tay River Jefferson lake Okotoks Nisku W-Pembina Sichuan Basin, SW China Early Triassic oolitic reservoirs Bohai Bay Basin

CO2 content (%)

H2S content (%)

21.3 5.3 40.7 11

5.5 15.2 5 12

10.8 3.5

6.2 6.2 31.3 10.1 11 0.1-34.5

34 34

0.5-9

8.8-17.1

not reported

40-92

(in the present context the term petroleum includes any type of organic matter, i.e., carbohydrates, kerogen, crude oil, bitumen, dissolved and gaseous organic compounds). Hence, redox reactions occur, either with or without the mediation of microbes. These reactions are usually discussed from the vantage point of sulfate reduction, which always implies concomitant hydrocarbon oxidation. Bacterially mediated sulfate reduction is called BSR (bacteri-

Figure 6.5.12. Thermal regimes of BSR and TSR

al sulfate reduction). Taxonomically, some of the involved microbes may not be bacteria but rather archeae or other types of microbes. Thus, strictly speaking, BSR should be called M(icrobial)SR. For historical reasons, however, the term BSR is retained. Abiological sulfate reduction is called TSR (thermochemical sulfate reduction). TSR is also called non-biogenic, non-microbial, organic, thermal, and abiological sulfate reduction.

Origin and distribution of non-hydrocarbon gases

The first report of BSR, although not named as such at the time, was by Cohn (1867), who discovered that the common sulfur bacterium Beggiatoa can form H2S under certain conditions. The group of sulfate-reducing bacteria was then identified and introduced by Beijerinck (1895) in a journal of bacteriology, parasitology, infectious diseases, and hygiene. BSR occurs in a large variety of sedimentary and low-temperature (less than about 80 °C) diagenetic environments (Fig. 6.5.12), which range from groundwater aquifers to marine sediments, reefal carbonates and layered or diapiric evaporites, and clastic sediments and rocks. TSR was introduced into the geoscientific literature by Toland (1960) who performed hydrothermal experiments using a variety of dissolved sulfates and hydrocarbons. Later experiments showed that TSR can take place at temperatures as low as 175 °C under laboratory conditions (summarised in Trudinger et al. 1985). Geological evidence suggests lower minimum temperatures of 100140 °C (Machel 1998a). Thus TSR is restricted to diagenetic and hydrothermal environments with temperatures in excess of about 100-140 °C (Fig. 6.5.12), and probably is the main process generating large quantities of hydrogen sulfide (up to about 90 vol-%) in numerous deep subsurface sour gas provinces of the world. In northwestern Europe TSR is associated with the carbonate and sulfate deposits of the Permian Zechstein. Considering the available evidence on the distribution of sour gas, burial depths and palaeotemperatures at the time of sulfate reduction, it is likely that most sour gas and metal sulfide deposits in the subsurface of northern Europe were formed by thermochemical sulfate reduction (Jochum 2000; MittagBrendel 2000). However, BSR was the only process generating H2S in at least some regions, such as northwestern Germany near the border to the Netherlands.

6.5.4.3 Important reaction steps Reaction paths and products for BSR and TSR have been summarised in several overview papers (e.g., Noeth 1997; Machel 2001, and references therein). The following discussion follows the reaction scheme proposed by Machel (2001), with minor modifications based on recent experimental studies by Zhang et al. (2007, 2008). All reaction steps have been demonstrated experimentally, observed in natural environments, calculated thermodynamically or inferred on the basis of elemental or isotopic evidence. Most redox steps are valid for both BSR and TSR. However, the systems considered here are assumed to be initially free of base and transition metals (the addition of these metals is discussed further below). Also, for simplicty, metastable or transient reaction products such as polysulfides that occur in subreactions are not included. In

reality, each reaction step consists of several sub-reactions and involves a net mass and charge balance. Aerobic or anaerobic biodegradation of non-gaseous hydrocarbons, i.e., crude oil, is important in the present context, as represented by reaction step 6.5.3: paraffinic hydrocarbons → biodegraded hydrocarbons (6.5.3) This process appears to be a prerequisite for BSR because sulfate-reducing bacteria depend on the metabolic residues of biodegradation as nutrients. When the system is anaerobic, sulfate-reducing bacteria can live symbiotically with methanogenic bacteria and utilise methane as a carbon source. Conversely, reaction 6.5.3 is not a prerequisite for TSR because abiologic sulfate reduction can take place with a variety of non-biodegraded hydrocarbons. Reaction step 6.5.4 represents thermal maturation of crude oil, which, among other reaction products, produces H2S, methane, and small amounts of polysulfides (PS): crude oil → light crude oil + H2S (+ PS) + CH4 (6.5.4) Reaction 6.5.4 may be important for TSR because H2S and/or its dissociation products (HS- and/or S2-, depending on pH) and/or the polysulfides appear to be catalysts for reaction steps 6.5.5 and 6.5.7 via reaction step 6.5.6. The initial S-O bond rupture and reduction of S6+ to lower valence states (most S6+ is reduced directly to S2-) is represented by several alternative reactions, depending on the types of reactive organic compounds and pH: 4R-CH3 + 3SO42- + 6H+ → 4R-COOH + 4H2O + 3H2S (6.5.5a) R-CH3 + 2R=CH2 + CH4 +3SO42- + 5H+ → 3R-COOH + HCO3- + 3H2O + 3H2S (6.5.5b) 2CH2O + SO42- → 2HCO3- + H2S (6.5.5c) In reaction 6.5.5a, the organic substrate compounds are high molecular-weight n-alkanes, in reaction 6.5.5b they are alkanes, alkenes, and methane, and in reaction 6.5.5c they are carbohydrates. Other reaction schemes could be formulated for different reactive hydrocarbon mixtures, or with methane as the only carbon source. Important reaction products are organic acids, bicarbonate ions and hydrogen sulfide (and/or, depending on pH, their associated and dissociated species, i.e., CO2, HS-, etc.). Possible reaction byproducts are solid bitumens from polymerisation. If methane is the only carbon source, no solid bitumen is formed. Some sulfur species with intermediate valence states, including polysulfides, may be formed during TSR (also omitted for simplicity).

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Elemental sulfur is commonly formed by partial reoxidation of H2S. Reaction steps 6.5.6a through 6.5.6d represent alternative ways for the generation of S0: 0

2H2S + O2 2S + 2H2O 3H2S +SO42- + 2H+ 4S0 + 4H2O H2S + SO42- +2H+ S0 + 2H2O +SO2 H2S + petroleum → S0 + altered petroleum S2- → S0 + 2 e-

(6.5.6a) (6.5.6b-1) (6.5.6b-2) (6.5.6c) (6.5.6d)

Reaction 6.5.6a represents abiological oxidation of H2S with molecular oxygen. This is likely to occur only where H2S escapes into the atmosphere or comes into contact with dissolved oxygen in hydrological systems. Reactions 6.5.6b-1 and 6.5.6b-2 are abiological oxidation reactions with excess dissolved sulfate at low pH. This type of sulfide oxidation is possible only abiologically at elevated temperatures. Water and elemental sulfur are the only reaction products when the H2S concentration is high (Reaction 6.5.6b-1), whereas sulfur dioxide is produced along with elemental sulfur where relatively little hydrogen sulfide is available (Reaction 6.5.6b-2). Reactions 6.5.6b-1 and 6.5.6b-2 are not the only reactions between sulfate and hydrogen sulfide. Others yield polysulfides, sulfite, and thiosulfate, which are known to be more reactive than sulfate in reaction 6.5.5 or elemental sulfur in reaction 6.5.7. Also, at elevated temperatures above about 100 °C, H2S may react abiologically with petroleum to produce elemental sulfur and NSO-compounds, if catalysts such as clays are present (Reaction 6.5.6c). Reaction 6.5.6d represents various microbial processes. Oxidation of sulfide to sulfur may be mediated by anaerobic or aerobic microbes. The electron acceptors during bacterial sulfide sulfur oxidation may be different for various genera and environments, and usually involve several enzyme-catalysed steps. Therefore, the simplest representation of the reaction paths during bacterial sulfide to sulfur oxidation is the valence change of sulfur (Reaction 6.5.6d). Regarding mass balance, however, reaction 6.5.6a may be taken to represent sulfide to sulfur oxidation in natural oxygenated environments, and reaction 6.5.6c may be taken to represent anaerobic bacterial sulfide to sulfur oxidation. Reaction 6.5.6b is probably not an alternative in this context, because microbes do not normally couple sulfide oxidation and sulfate reduction. Most of the elemental sulfur formed in reaction(s) 6.5.6 may not survive as a separate phase in a natural environment. If the partial pressure of oxygen is high enough and/ or certain types of microbes are present, H2S may be oxidised inorganically or microbially to SO42-. Conversely, if the environment is sufficiently reducing, reduction of S0 to

S2- takes place, as represented by the following simplified redox reaction step: 4S0 + 1.33 (-CH2-) + 2.66 H2O + 1.33 OH- → 4H2S + 1.33 HCO3 (6.5.7) Reduction of S0 may take place because sulfur-reducing bacteria may be present, or abiologically at temperatures in excess of about 100 °C because S0 is an active oxidising agent for hydrocarbons, symbolised by the functional methylene group (-CH2-). Main reaction products are bicarbonate ions and hydrogen sulfide (and/or CO2 and HS-, as in the case of reaction step 6.5.5c). In natural environments with crude oil as the carbon source, reaction 6.5.7 also generates naphthenic acids, aromatic compounds, bitumen, and numerous inorganic and organic sulfur compounds (i.e., polysulfides, mercaptans, thioalkanes, thiophenes). Reactions 6.5.5, 6.5.6, and 6.5.7 produce compounds that are partially recycled (i.e., H2S, S0, polysulfides) and therefore do not appear explicitly in the net mass balance: petroleum + SO42- → altered petroleum + solid bitumen + HCO3- (+CO2) + H2S (+HS-) + H2O (6.5.8) Polysulfides and other compounds with intermediate valence states of sulfur, although undoubtedly present, are not included in reaction 6.5.8 for simplicity. Besides the inorganic products (CO2, H2O, H2S), appreciable amounts of NSO-compounds are formed (collectively called “altered petroleum”). In diagenetic environments, reactions 6.5.3 through 6.5.8 may or may not take place, and they may proceed in a different order, depending on the presence or absence of certain microbes and hydrogen sulfide. During BSR, complex associations of microbes that are different for each redox step perform multiple intermediate reaction substeps. Initially, the sequence of reactions would probably be: Reaction 6.5.3 and then 6.5.5. If the environment contains some oxygen (i.e., at the groundwater/oil interface, at an oil/air interface, or in atmospheric oil seeps), reactions 6.5.6a (or 6.5.6c, 6.5.6d) will proceed to form S0. This sulfur will accumulate as a net reaction product only if the microbes mediating reaction 6.5.7 are not present, or if the environment is not conducive to growth, for example, if “digestible” hydrocarbons are no longer available. Once instigated, reactions 6.5.5, 6.5.6, and 6.5.7 proceed simultaneously at or near the aerobic/anaerobic interface, because some, if not most, microbes promoting reaction 6.5.6 are aerobic, whereas those promoting reactions 6.5.5 and 6.5.7 are anaerobic. This interface could be an oil/water contact or a diffusion zone within anaerobic sediments overlain by aerobic water. Should the whole

Origin and distribution of non-hydrocarbon gases

environment become closed and totally anaerobic, reactions 6.5.5 and 6.5.7 will take over until the nutrients and reactants are depleted. In most natural TSR settings, the initial sequence of reactions is probably 6.5.4 (providing the initial hydrogen sulfide and/or polysulfides for reaction 6.5.6, then 6.5.6b or 6.5.6c (also forming polysulfides, sulfite, thiosulfate), then reactions 6.5.5 and 6.5.7. Once instigated, these reactions proceed simultaneously, because H2S and polysulfides from reactions 6.5.5 and 6.5.7 are partially recycled in reaction 6.5.6. As in the case of BSR, S0 accumulates as a net reaction product only if reaction 6.5.7 does not proceed or is retarded. In laboratory experiments, TSR has also been instigated directly via reaction(s) 6.5.5, i.e., in the absence of metastable organic S-compounds (Zhang et al. 2007). Also, further experimental studies and calculations suggest that dissolved bisulfate (HSO4-) probably is much more reactive than dissolved sulfate (SO42-), possibly involving MgSO4 ion pairs (Ma et al. in press). These types of studies show that the above reaction scheme for BSR-TSR, although correct in principle (as it properly describes the major reaction products and by-products), is necessarily crude and incorrect in detail.

6.5.4.4 Common Products in Natural Environments The most common reaction products of BSR and/or TSR are H2S, calcite and/or dolomite, elemental sulfur, and solid bitumen (Fig. 6.5.13). This paragenesis has been found in many BSR and TSR settings and can be considered indicative of sulfate-hydrocarbon redox-reactions in general. A distinction of BSR versus TSR commonly requires additional geochemical data, however, as discussed further below.

Figure 6.5.13. Core sample from a sour gas field in the Devonian Southesk Cairn Carbonate Complex, Alberta, Canada: intergrowth of calcite (white), elemental sulfur (yellow), and solid bitumen (black). This paragenesis is a product of thermochemical sulfate reduction and common in deeply buried carbonate rocks (approximately 3000-4500 m) that contain(ed) anhydrite and oil or gas condensate. The major gaseous reaction product is hydrogen sulfide. Field of view: ~3 cm

During BSR, H2S can be generated as long as reactants and nutrients for the microbes are available, and as long as the H2S concentration is below the toxic level for the microbial metabolism. Hence, for large quantities of H2S to be formed, the system must be open and allow for continuous inward diffusion of sulfate if the system contains hydrocarbons, or of hydrocarbons if the system contains sulfate. The evolving H2S is fixed as metal sulfides, in organic compounds, or may escape as gas from the reaction site, and thus may not accumulate to form a reservoir. However, the possibility of biogenic H2S migrating away and forming a sour gas reservoir elsewhere (i.e., away from the actual site of BSR) cannot be ruled out.

kerogens can also liberate sour gas in the absence of TSR. Using crude oil with about 3% sulfur as representative of sulfur-rich crude, Orr (1977) calculated that no more than 2-3 vol-% H2S (in reservoir gas) can evolve from thermal cracking of crude oil. Some kerogens, such as the Monterrey shale, have up to about 20% sulfur and thus may generate gas with higher H2S contents. Deep reservoirs with 20-90 % H2S are invariably due to TSR.

On the other hand, deep reservoirs with more than a few percent of H2S are always due to TSR, although sulfur-rich

Elemental sulfur accumulates as a net reaction product if reaction step 6.5.7 is inhibited. This happens if the system runs out of reactive hydrocarbons (BSR/TSR), or if reactive hydrocarbons are not supplied fast enough (TSR). BSR and TSR form many distinctive NSO-compounds (e.g., Ho et al. 1974; Manzano et al. 1997; Zhang et al 2007, 2008; Kelemen et al. 2008). Some of these compounds are contained in the solid organic precipitates

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tianite) are also formed as a result of BSR/TSR where the respective metals are available. If transition or base metals are present, disseminated or stratiform base metal or Mississippi-Valley-Type deposits may form. Minerals such as pyrite, galena, and sphalerite are precipitated as a result of sulfide generation during reactions 6.5.5 and 6.5.7. In these cases, partial dissolution of the host rock should occur (at least initially), because the precipitation of sulfides generally is associated with strongly acid-generating reactions. In addition, BSR/TSR themselves could generate acidity due to the release of CO2, perhaps in reaction 6.5.5. However, the acidity resulting from this carbon dioxide is probably low and negligable in most natural environments. On the other hand, much larger quantities of CO2 evolve from thermal cracking of kerogen or crude oil. Hence, dissolution by thermogenic carbonic acid in deeply buried hydrocarbon reservoirs may compound potential dissolution caused by precipitation of sulfides.

Figure 6.5.14. The difference in sulfur isotope composition (CDT = Canyon Diablo Troilite Standard) between sulfate dissolved in formation waters and hydrogen sulfide versus depth from three sedimentary basins (after Krouse 1977). High isotope difference values at depths less than about 2000 m indicate in situ bacterial sulfate reduction, the effect of which diminished with increasing depth. At depths greater than about 2200 m, bacterial sulfate reduction appears to be volumetrically insignificant and all H2S is thermogenic. Figure reproduced from Machel and Foght (2000)

(bitumen), which should be formed where hydrocarbons other than methane are utilised. Others are contained in the gas phase. These compounds can be distinguished from those generated by biodegradation or thermal maturation under favorable circumstances. Several minerals may precipitate as the result of BSR/ TSR if the respective metal ions are present or are transported to the redox reaction site, or if the reaction products of BSR/TSR are transported into an environment containing metal ions. Firstly, the presence of alkali earth metals will result in the precipitation of carbonates (mainly calcite and dolomite), either as a cement or as a replacement of the dissolving sulfates (mainly gypsum and anhydrite). This is a direct result of reaction steps 6.5.5 and 6.5.7, which generate bicarbonate. Additionally, the reaction of polysulfides with bicarbonate may also lead to calcite precipitation, and bicarbonate may form as a result of bacterial oxidation of sulfide to sulfur (reaction step 6.5.6), with subsequent precipitation of this bicarbonate as calcite. Other carbonates (ankerite, siderite, witherite, stron-

Subordinate reaction by-products that may form in the vicinity of BSR and TSR as a result of released and/or consumed phases are cerrusite, barite, fluorite, and gases such as nitrogen, and helium. The availability and mass proportions of the reactants determine which products and by-products are formed, and in which proportions. The net mass balance reaction 6.5.8 also contains water as a reaction product. There is evidence for the release of water during TSR in some TSR settings, such as the Khuff Formation (Worden et al. 1996), but water does not appear to be a volumetrically significant by-product of TSR in many other TSR settings (Machel 1998b). TSR is likely to take place in the gas-water transition zones, and geochemical anomalies are confined to these zones, or even to individual TSR-reaction sites within such zones. TSR can occur throughout the entire gas leg only in a reservoir without a clear gas-water transition zone, and where the gas zone contains significant amounts of water to facilitate dissolution of anhydrite.

6.5.4.5 Temperature ranges and reaction kinetics of BSR and TSR The major reaction pathways are similar for biological and abiological sulfate-hydrocarbon redox reactions. However, BSR and TSR appear to be mutually exclusive processes in natural environments. Among the best indications for this phenomenon are the natural occurrences and isotopic compositions of H2S, suggesting that BSR and TSR take place at two particular temperature/thermal maturity levels: at low levels of less than about 80 °C (equivalent to about 0.2-0.3% VRr), and at high levels in

Origin and distribution of non-hydrocarbon gases

excess of 100-140 °C (>1.5% VRr) (Fig. 6.5.12). These temperature-maturation correlations are only generalised approximations because of the time-dependence of thermal maturation; see Littke et al. this volume). The lower temperature/maturity level coincides with BSR because almost no sulfate-reducing bacteria can thrive at the temperatures of the upper level. Most geologically significant microbes, i.e., sulfate reducers and methanogens, cease metabolism at about 80 °C (Machel and Foght 2000), while most sulfate-reducing bacteria do not metabolise at temperatures in excess of about 45 °C (at 1 atm). Hence, with odd exceptions, BSR takes place in near-surface and shallow subsurface environments. In these environments, BSR proceeds almost instantaneously on a geological time scale and does not depend on any catalysts (other than enzymes within the cells). This has been demonstrated experimentally in the laboratory, and in the field. For example, water injection for enhanced petroleum recovery in several shallow Russian and Canadian oil fields resulted in rapid H2S production. Within less than 10 years, up to about 10% H2S was generated in previously sweet reservoirs because of the introduction of bacteria into hitherto sterile reservoirs. Because of the limited temperature tolerance of most microbes, BSR has a relatively shallow depth limit in natural environments. The maximum depth for BSR is

probably close to 2000 m, as shown by isotope data (Fig. 6.5.14). This depth is not the depth limit of the biosphere, however, as other types of microbes have been identified much deeper. For example, Ashirov (1962) and Rosanova and Khudyakova (1974) reported microbes from 3290 m (equivalent to about 75 °C and a pressure of about 350 bars). In such deep environments large-scale BSR is essentially prohibited by the inability of most sulfate reducers to metabolise and/or by exceedingly slow reaction rates. Hence, sulfate reduction, the associated biodegradation of hydrocarbons, and aerobic sulfide to sulfur oxidation, are generally bracketed between the “aeration zone” of exposed rocks (0-200 m depth) and the bottom of the oxygenated groundwater zone (generally about 600 m). Accordingly, carbonates, bitumen, and elemental sulfur resulting from BSR and partial re-oxidation of sulfide occur generally in near-surface environments (i.e., in salt dome cap rocks of the United States and Europe: (e.g., Feely and Kulp 1957) and along/below the oil-water interface in shallow oil reservoirs. Meteoric water is probably the vehicle that brings microbes into contact with subsurface hydrocarbon pools, i.e., via faults, fractures, and other conduits, and oxygenated groundwater is probably the cause of H2S oxidation to S0 in most shallow sulfur deposits. Where solid bitumen is absent, and the stable carbon isotope ratios of precipitated carbonates are lower than those of oil, methane is indicated as the main reactant hydrocarbon.

Table 6.5.4. Sulfur-, oxygen- and carbon isotope fractionations during BSR and TSR (after Machel et al. 1995b) δ34S fractionation of inorganic compounds SO42- -> S2-:

BSR:

-15 to - 65 ‰

maturation:

TSR:

-20 ‰ @ 100 oC -15 ‰ @ 150 oC -10 ‰ @ 200 oC <-2 ‰

S2- -> S0:

oxygen: sulfate: hydrocarbons: anaerobic bacteria: aerobic bacteria:

-5 to -7.5(?)‰ +/-0 to -3.7‰ ? +/-0 to +2‰ +/-0 to -18(?)‰

S0 -> SO42-:

+/-0‰

δ13C fractionation of carbonates BSR /TSR

CH4 oxidation: solid/liquid hydrocarbon oxidation:

up to -70‰ PDB -20 to -30‰ PDB



carbonates depleted relative to marine carbonate:

up to -50‰ PDB

δ18O fractionations of carbonates BSR /TSR carbonates depleted relative to marine carbonate:

up to - 8 ‰ PDB

453

454

B.M. Krooss · B. Plessen · H.G. Machel · V. Lüders · R. Littke

Figure 6.5.15. Distribution of H2S in Permian Zechstein gas fields of northern Germany. The fields in the contoured region south of Bremen contain up to 35% H2S from TSR. The smaller fields west of the Ems river were not exposed to sufficiently high temperatures for TSR. Their H2S contents of up to 2% are likely to result from BSR. After Mittag-Brendel (2000)

TSR appears to take place only at temperatures >100140 °C, >1.5 VRr (Fig. 6.5.12), and is much slower than BSR, yet TSR is still fairly fast geologically and probably goes to completion in less than one million years at any given site (Machel 2001). The activation energies are relatively large and differ for various hydrocarbons (experimentally shown: e.g., Toland 1960; Zhang et al. 2007, 2008). Accordingly, TSR has been performed in the laboratory only at high temperatures (in excess of 175 °C), and geologically significant reaction rates were measured only above about 250 °C. However, there is ample geological, theoretical, and circumstantial evidence for the occurrence of TSR at diagenetic temperatures as low as at least 135-140 °C (summarised in Machel 1998a, 2001). As a generalisation, there is a range of minimum temperatures, and the minimum temperature for TSR may be variable from pool to pool even within a single sour gas play (as in the Jurassic Smackover Formation of the US Gulf Coast region: Claypool and Mancini 1989); in some locations there is a specific minimum temperature for TSR represented by a certain minimum depth, as in the Khuff Formation (Worden et al. 1995) and in the Nisku Formation (Machel et al. 1995b). The lowest confirmed temperature for TSR to date is 127 °C, with questionable evidence pointing to a temperature as low as perhaps 100 °C.

There are a number of reasons why TSR should have a range of possible minimum temperatures, rather than a sharp minimum temperature. The governing factors include the composition of the reactive hydrocarbons (including their sulfur contents), catalysts, anhydrite distribution (availability) and dissolution rate, wettability and migration/diffusion rates of the major reactants, all of which may vary from region to region and even from pool to pool. For example, the very fact that different types of hydrocarbons with different bond strengths and TSR activation energies are involved necessitates that there is not just one TSR minimum temperature but that this temperature is a function of the availability of reactive hydrocarbons. Isotopic and/or gas-chromatographic studies have identified branched chain and n-alkanes in the gasoline range as the most reactive hydrocarbons, followed by cyclic and monoaromatic species in the gasoline range, with methane being the most stable and least reactive “ The amounts and types of organic sulfur compounds can dramatically influence the onset of TSR (Reaction step 6.5.4). Wettability and water saturation are other commonly overlooked yet important factors in this context. Just as solid sulfate cannot react at typical TSR reservoir temperatures, TSR cannot occur in oil-wet (gas-wet)

Origin and distribution of non-hydrocarbon gases

reservoirs because sulfate cannot dissolve. Hence, variations in wettability alone could account for at least some of the observed temperature “cross-overs”, such as those in the Smackover trend, that is, some reservoirs that are hot enough for TSR may be inhibited from H2S formation simply because the water phase is not capable of dissolving anhydrite (or not fast enough). “Cross-overs” at the reservoir scale may well be the result of variations of wettability from pool to pool (or from low water saturation). Even in systems that are water-wet, low water saturation can effectively inhibit TSR because the water would be effectively immobile and sulfate-hydrocarbon interaction would be greatly inhibited. Conversely, a well-defined, specific minimum temperature for TSR can be expected only where the reservoir conditions are fairly homogeneous on the scale of a field or a play. Also, variations in wettability can account for variations in TSR phenomena on the smaller scale of individual TSR reaction sites. An example would be the small-scale variations and trends in δ18O, δ13C, and fluid inclusion data of TSR calcites and TSR dolomites found within individual sour gas pools of various locations.

6.5.4.6 Heat released Early studies have asserted that BSR and/or TSR are exothermic processes and estimated the heat released as about 30 - 40 kcal/mol hydrocarbon (Bush 1972), or 10 kcal/mol calcium sulfate (Dhannoun and Fyfe 1972). These studies suggested or inferred that “hot spots” could be generated via reactions between sulfate and hydrocarbons. Indeed, some sour gas reservoirs are warmer than the surrounding rock units by up to about 10 °C. However, sour gas reservoirs are rarely significantly (more than 10 °C) warmer than the surrounding rock units. Moreover, it is now known that the net reaction of TSR can be exothermic or endothermic, depending mainly on the type(s) of organic reactant(s) and/or net reaction stoichiometry (Simpson et al. 1996). The question arises whether observed heat anomalies in association with sour gas reservoirs are the result of TSR or coincidental occurrences. This question can be approached from the vantage points of reaction kinetics and of heat conduction and convection. As indicated in the previous section, TSR appears to be a relatively fast process. Hence, if a sour gas set-

Table 6.5.5. δ34S and δ13C variations of organic compounds (after Machel et al. 1995a) δ34S variations of NSO-compounds depleted ca.-15 ‰ (relative to SO42-) unchanged relative to immmature kerogen depleted ca.-15 ‰ (relative to SO42-) <+2 ‰ enriched (relative to kerogen) enriched or depleted (relative to kerogen and crude) enriched or depleted (relative to kerogen and crude) slightly depleted (relative to kerogen and crude) strongly enriched (relative to kerogen and crude)

Kerogen (immature): Kerogen (mature): Crude (immature): Crude (mature): Crude BSR: Solid bitumen BSR: Solid bitumen thermal: Solid bitumen TSR: δ13C variations of organic compounds Kerogen Crude (immature): Crude (mature): Crude Crude Solid bitumen Condensate

--maturation-->

--maturation-->

heavier (up to 3 ‰) lighter (ca. 3 ‰) relative to kerogen unchanged relative to kerogen heavier (thermal cracking)

--maturation-->

lighter (gas deasphalting)

--maturation -->

enriched up to 3 ‰ (relative to light crude) unchanged (relative to heavy crude) same as solid bitumen from maturation depleted (up to 7‰) relative to solid bitumen from maturation

--maturation --> --BSR--> --TSR-->

--TSR-->

enriched (up to several ‰)

455

456

B.M. Krooss · B. Plessen · H.G. Machel · V. Lüders · R. Littke

Table 6.5.6. Differentiation between fluids derived from maturation or BSR and TSR (after Machel et al. 2005a) Solid bitumens

formed by maturation or BSR

formed by TSR

often soluble

partly to largely insoluble

ca. -15‰

-15 to 0‰

rel. high 4 to 10 relatively low absent

rel. low 10 to 28 positive correlation up to 65% present

increasing

decreasing?



TSR observed deviation from No-TSR condensate

TSR observed total range

S % δ13C [CH4] ‰ PDB δ13C [C2H6] ‰ PDB δ13C [C3H8] ‰ PDB δ34S ‰ CDT sat/arom C15+ % gas conversion % °API CO2 % H2S %

+0.8 to +1.3 -1.4 to -2.9 up to +10 up to +10 +5.0 to +23.7 -5.1 to -8.6 [not always] -9.3 to -29.1 +11 to +37 -3.1 to -9.3 +0 to +10

+0.1 to +1.4 - 46 to -35 - 26 to -18 - 22 to -18 +2.4 to +24.1 +0.2 to +8.8

up to 95

up to 95

solubility δ34S ke/cr-bit (rel. to SO42-) H/C S % δ34S vs. S % asphaltenes % So δ13C vs. depth of sat, NSO, asph Gas condensates

ting is warmer than the surrounding rock units, TSR may be actively occurring, or the heat liberated earlier during TSR may have not yet dissipated. The latter possibility can be quantified with standard methodology from igneous petrology. Temperature anomalies caused by much stronger heat sources, such as igneous intrusions, usually dissipate in geologically short periods of time. The time for heat to dissipate by conduction alone can be calculated. A heat anomaly of perhaps 5 to 10 °C could originate from TSR if the reactions are exothermic and while TSR is taking place, but the small amounts of heat that could be generated quickly dissipate. Alternatively, a measured temperature anomaly could be unrelated to TSR but result from hydrodynamics (hot formation waters flowing upwards relatively rapidly) or increased heat flow from the basement.

6.5.4.7 Distinguishing between BSR and TSR As discussed in the previous sections, BSR and TSR may form the same organic and inorganic reaction products.

+11.3 to +40.4 +48 to +85 +42.3 to +56.0 2.4 to 50.1

Therefore, geochemical criteria are necessary to distinguish between BSR and TSR where the geological evidence is insufficient for this discrimination. Also, some criteria are necessary to distinguish reaction products of BSR and/or TSR from other processes that may form the same or similar products. For example, solid bitumen can also form during simple thermal maturation of hydrocarbons, from inspissation, or from aerobic biodegradation. Similarly, carbonate and sulfide precipitation may take place without any involvement of sulfate-hydrocarbon redox reactions. These aspects are discussed in Machel et al. (1995a), Zhang et al. (2007, 2008), and Kelemen et al. (2008; see Tables 6.5.4, 6.5.5, 6.5.6).

6.5.4.8 H2S-rich gases in the North German Basin In the German part of the Central European Basin System natural gases with H2S occur in Permian Zechstein reservoirs (Stassfurt Carbonate; Table 6.5.3; Fig. 6.5.15). The H2S contents of up to 35% in the region between the Bramsche Massif and Bremen, which had high heat

Origin and distribution of non-hydrocarbon gases

Figure 6.5.16. P-T diagram showing trapping conditions of primary and co-genetic aqueous and CH4-CO2 and C-H-O-S±N inclusions in minerals from selected samples from the Lower Saxony Basin via isochore construction. Isochores were calculated with the MacFlinCor computer program (Brown and Hagemann 1994). P-T conditions of fluid entrapment are shown as dark blue fields for CH4-CO2 inclusions and the pink field for H2S rich inclusions hosted in hydrothermal minerals from Zechstein wells in the Lower Saxony Basin (modified after Reutel et al. 1995)

flow in the past, are the result of TSR (Reutel et al. 1995; Mittag-Brendel 2000). H2S also occurs in Permian reservoirs in the northwestern part of Germany near the German-Netherlands border, but only up to about 2% (Fig. 6.5.15). These reservoirs are located at relatively shallow depths and never reached the minimum temperatures necessary for TSR. Based on compositional and isotopic data such as those presented in the previous section, H2S in these reservoirs is due to BSR (Mittag-Brendel 2000). Specifically, methane-ethane-propane ratios and their δ13C isotope ratios, when correlated with increasing H2S contents, permit a positive identification of TSR versus BSR in these two regions. Also, the δ34S ratios of BSR-generated H2S ranges from about –10 to +10 permil CDT (Canyon Diablo Troilite Standard), as would be expected from a microbial system with Rayleigh fractionation. Conversely, the δ34S ratios of TSR-generated H2S approach that of the Zechstein anhydrite with increasing H2S content, as would be expected from a TSR system (Tables 6.5.4 and 6.5.5). Furthermore, the BSR gases have highly variable CO2 contents that range from near zero to about 50%, while the TSR gases display a very narrow range of about 7-10% CO2. The CO2 in the BSR reservoirs is enriched in 13C, which points to microbial methanogenesis as a key factor in the generation of this CO2 (the carbon isotopic fractionation between methane and carbon dioxide is very large and provides for 13C-de-

pleted methane associated with 13C-enriched carbon dioxide). Thus, both the H2S and significant amounts of CO2 in these reservoirs is microbial. On the other hand, the CO2 in the thermogenic gases is a by-product of TSR, as are the associated calcite and elemental sulfur that have been identified in core samples

6.5.5 Evidence from vein mineralisation and fluid inclusions Most carbonate reservoir rocks host numerous fissure or even vein mineralisations, which contain recrystallised carbonates and/or late diagenetic hydrothermal carbonates, fluorite, or sulfates. These minerals have been studied in great detail microthermometrically (e.g., Zwart 1995; Mittelstädt 1992; Reutel and Lüders 1998; Lüders et al. 2005) as well as using the stable isotopic (C-O) composition of carbonate host rocks and authigenic calcites (Fischer et al. 2006). The studies show that the late-stage hydrothermal minerals formed at temperatures of about 150°C from highly saline CaCl2-NaCl-H2O brines. Mineral formation is assumed to have occurred from Early to Middle Jurassic through Middle Cretaceous times (Zwart 1995). Gaseous fluid inclusions trapped next to aqueous brine-type inclusions are either primary or secondary in

457

458

B.M. Krooss · B. Plessen · H.G. Machel · V. Lüders · R. Littke

Figure 6.5.17. Evolution of the diagenesis of Zechstein carbonates based on results of microthermometric studies. Path A represents wells in the central part of the Lower Saxony Basin (Rieken 1988). Path B is related to stages of pressure decrease related to inversion tectonics during the Early Cretaceous. The formation of H2S-rich gas reservoirs seems to be related to this stage of basin evolution. Locally, sour gas reservoirs were diluted by CH4 (Path C)

origin and show regional variations in composition. Zechstein carbonate samples from wells in the southern part of the Lower Saxony Basin contain CO2-rich fluids or CH4/ CO2 mixtures in fluid inclusions hosted by hydrothermal calcite and/or fluorite. In the northern part of the basin the composition of gaseous inclusions in the same minerals can be very complex. They may either contain pure CH4, mixtures of CH4 and CO2, or show variable C-H-O-S±N composition (Reutel et al. 1995). The trapping conditions of the gas-rich inclusions differed over time: C-H-O-S±N inclusions show phase transitions typical of low-density gas inclusions, indicating pressure conditions close to hydrostatic during fluid entrapment. In contrast, CH4-rich inclusions and/or CH4-CO2 mixtures were trapped under considerably higher pressure conditions (Fig. 6.5.16). The calculated pressure conditions for fluid entrapment of H2S-rich and CH4-CO2-bearing fluid inclusions suggest at least two stages of gas migration in the Low-

er Saxony Basin: C-H-O-S±N-bearing fluid inclusions preferentially occur in fluorite, which is younger than recrystallised calcite within the Stassfurt carbonate. CH4-CO2 or even pure CH4 inclusions occur along trails of secondary inclusions in fluorite or as primary inclusions in coelestine, which is always younger than fluorite, thus indicating that the formation of H2S-bearing gases is related to an earlier gas migration event (Fig. 6.5.17). The low pressure condition of entrapment of C-H-O-S±N inclusions suggests that these fluids were trapped during stages of tectonic uplift, most probably during the main stage of inversion during the Early Cretaceous. H2S-rich gas was formed by TSR at temperatures of about 150°C as a result of the interaction between CH4 and/or hydrocarbons with SO4--bearing formation water. In areas with subsequent burial, the migration of CH4 and/or CH4-CO2 gas mixtures occurred under increasing pressure conditions and probably led to a dilution of formerly H2S-rich reservoirs (Fig. 6.5.17).

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Shell International Petroleum Maatschappij, The Hague, Amsterdam Ziegler PA (1987) Late Cretaceous and Cenozoic intra-plate compressional deformations in the Alpine foreland. Tectonophysics 137:389-420 Ziegler PA (1988) Evolution of the Arctic-North Atlantic and the Western Tethys- A visual presentation of a series of palaeotectonic maps. Evolution of the Arctic-North Atlantic and the Western Tethys. AAPG Mem 43:164-196 Ziegler PA (1989) Evolution of the North Atlantic – an overview. In: Tankard AJ, Balkwill HR (eds) Extensional tectonics and stratigraphy of the North Atlantic region. AAPG Mem 46:111-129 Ziegler PA (1990) Geological Atlas of Western and Central Europe. Shell International Petroleum Maatschappij, The Hague, Amsterdam Ziegler PA, Dèzes P (2006) Crustal evolution of Western and Central Europe. In: Gee DG, Stephenson RA (eds) European Lithosphere Dynamics. Geol Soc Lond Mem 32:43-56 Ziegler PA, van Hoorn B (1989) Evolution of the North Sea Rift. In: Tankard AJ and Balkwill HR (eds) Extensional Tectonics and Stratigraphy of the North Atlantic Margins. Am Assoc Petrol Geol Mem 46:471-500 Ziegler PA, Cloetingh S, van Wees JD (1995) Dynamics of intraplate compressional intraplate deformation: the alpine forelandand other examples. Tectonophysics 252:7-22 Ziegler PA, van Wees J-D, Cloetingh SAPL (1998) Mechanical controls on collision-related compressional intraplate deformation. Tectonophysics 300(1-4):103-129 Ziegler PM (1975) Geological evolution of the North Sea and its tectonic framework. AAPG Bull 59:1073-1097 Ziljstra H (1995) The sedimentology of chalk. Lecture Notes in Earth Sci 54, Springer, New York, Heidelberg, Berlin Zimmermann H (2000) Tertiary seawater chemistry - Implications from primary fluid inclusions in marine halite. Amer J Sci 300:723-767

References Zirngast M (1996) The development of the Gorleben salt dome (Northwest Germany) based on quantitative analysis of peripheral sinks. Geol Soc Spec Publ 100:203-226 Zoback MD, Moos D, Mastin L, Andersson R Wellbore (1985) Breakouts and in situ stress. J Geophys Res 90:5523-5530 Zöllner H, Reicherter K, Schikowsky P (2007) Mesozoic and Cenozoic structural framework of the Bay of Kiel area, western Baltic Sea. Int J Earth Sci, doi:10.1007/s00531-007-0277-9 Zoubtsov S, Renard F, Gratier JP, Guiguet R, Dysthe DK, Traskine V (2004) Experimental pressure solution creep of polymineralic aggregates. Tectonophysics 385:45-57 Zühlke R, Bechstädt T, Mundil R (2003) Sub-Milankovitch and Milankovitch forcing on a model Mesozoic carbonate platform – the Latemar (Middle Triassic, Italy). Terra Nova 15:69-80 Zwach C, Throndsen T, Bergan M (2000) Quick mapping of basin modelling results - a key for quantifying prospect sensitivities. In: Ofstad K, Kittilsen JE, Alexander-Marrack P (eds) Improving the Exploration Process by Learning from the Past. NPF Spec Publ 9, Amsterdam Zwart EW (1995) Fluid inclusions in carbonate rocks and calcite cements. Ph.D. thesis, University Amsterdam Zwingmann H, Clauer N, Gaupp R (1998) Timing of fluid flow in a sandstone reservoir of the North-German Rotliegend (Permian) by KAr dating of related hydrothermal illite. In: Parnell J (ed) Dating and duration of fluid flow and fluid-rock interaction. Geol Soc Lond Spec Publ 144:91-106 Zwingmann H, Clauer N, Gaupp R (1999) Structure-related geochemical (REE) and isotopic (K-Ar, Rb-Sr, ∂18O) characteristics of clay minerals from Rotliegend sandstone reservoirs (Permian, northern Germany). Geochim Cosmochim Acta 63:2805-2823

505

Subject Index

3D restoration 100 3D seismic data 107, 115, 117-119, 190, 292, 293, 302 A Aalenian 166, 167, 181, 200ff Accomondation space 44, 219 Accretion 6, 22, 37, 41, 93, 94, 254, 258, 259, 434 Accumulation 3, 8, 10, 26ff, 37, 38, 115, 161, 165, 169, 188ff, 212ff, 251, 261ff, 302ff, 324ff, 359, 364, 365, 372ff, 411ff, 433, Accuracy 98, 327, 407 Across-fault flow 381, 383 Activity coefficient calculation methods 406 Adiabatry 258 Aegean 6, 41, 50 Aeolian 8, 163, 183ff, 189, 218, 259, 403, 408, 409 Africa 4, 25, 50, 56, 119, 125, 161, 181, 188, 206, 210, 256ff, 291, 413, 421 African plate 161 Age(s) Cooling ages 151, 168 Age Dating 111, 397 Albian 159, 167, 207ff, 220, 223, 318 Albitization Aller Lineament 309, 313, 318 Alluvial 162, 163, 171, 179, 184, 192, 232 Alpine deformation 94, 292 orogen 25, 58 Tethys 130 Alteration Rim 111, 266 Altmark (group) Basin 220, 221, 313 tectonic movements 176 Ambiguitiy Ammonium 401, 434, 435, 438, 440ff Amorican Massif 169 Analogue models 98, 107 Anatolia 5 Anglo-Dutch basin 301, 306 Anhydrite 133, 135, 169, 185, 188, 194, 205, 241, 245, 249ff, 303, 328, 354, 355, 358, 372, 380, 390ff 444ff Anisian 164, 191, 193 Anisotropy 46, 72, 334, 370, 371 Anomaly Bouguer 26, 87, 90, 93 Free air 26, 87-90 Pritzwalk 86, 94 residual 84, 87, 89, 90 Anthracite 138, 139, 360, 435, 439, 440 Anticline107, 113, 114, 213, 214, 226, 228, 229, 269, 299 Apatite 142, 150, 151, 186, 228, 229, 391, 397, 398 Apatite – fission tracks 142, 149 Aquifer 9, 264, 331ff, 351ff, 386, 387, 408, 410, 449 Ara Group South Oman 330 Ara Salt 261, 329, 330, 430, 431 Arabia 25, 259, 261

Arctic Ocean 125, 128, 159, 184, 186, 200, 210 Arid 28, 157, 159, 170, 171, 177, 182, 187, 191, 194, 195, 199, 210, 250ff, 305, 399 Armorica 41 A seismic 43 Asthenosphere 4, 19, 38, 39, 41, 42, 45, 48, 60, 63, 66, 80, 82, 95, 235 Asthenospheric Mantle 235 Atacama Desert 256 Atlantic 41, 49, 50, 59, 125, 130, 161, 166, 168, 184ff, 235, 254ff, 291, 292, 305 Atlantic rift 159, 192, 197, 205, 206, 209, 261 Attraction 86 Aulacogene 5 Austro-Carpathians 130, Authigenic 391-393, 395-398, 400, 401, 403, 404, 409, 443, 446, 457 Autunian 43 Avalonia Baltica Suture 67, 79, 82, 83, 89, 90 B BABEL 19, 57, 61, 79, 81 Back arc 5-7, 37, 416 Back-stripping 161 Backthrust 6, 48, 113 Baikal lake 38 Baikal Rift 4 Bajocian 166, 167, 200ff Baltic Sea 19, 68, 76, 77ff, 94, 120, 233, 236, 238, 244, 245, 320, 321, 337 Baltic Shield 19, 20, 81, 82 Baltica 19, 20, 22, 67, 76, 77, 79, 80, 82, 83, 85, 89, 90, 93, 94, 95 Baltic-East European craton 19, 20, 22, 60, 61, 83, 87, 95, 125, 143, 210 Baruth 233 Basal Detachment 233 Basaltic, basalt 4, 24, 162, 173-176, 178, 184, 201, 206, 266, 356, 369, 371, 435, 436 Base level 161, 162, 166, 169, 212ff Basement crystalline 82, 89 fault 48, 59, 71, 94, 104ff, 128, 198, 214, 234, 235, 237, 268, 272, 297ff, 308, 312, 313 folding 110, 113, 212, 213, 215, 220 uplift 104, 105, 109-111, 114, 219, 221, 223 Basin almost-filled 5 back arc 5-7, 37, 416 Barnim 178, 179 Collio 173, 175 compound 5, 13 craton interior 9 Danish 24, 57, 66ff, 87ff, 125ff, 171, 182ff, 228, 231, 308ff differentiation 152, 170, 200, 220, 308, 312 evolution 3ff, 29, 38ff, 94, 104, 125, 129, 157, 173, 226, 277, 278, 292, 305, 909, 389, 401, 412, 442, 458 foreland 22, 97, 166, 181, 211, 212, 227, 261, 307, 425

508

Subject Index formation 5, 6, 22, 37ff, 81, 181, 211, 215 Illinois 5 initiation 74, 126, 173ff intracratonic 37, 265, 269 intra-mountain 5 intra-Sudetic 173 inversion 7, 60, 109, 125, 184, 211ff, 268, 306, 308, 315 Michigan 5, 260ff Modelling 37, 52, 98, 125, 134, 224, 367, 386, 412ff Molasse 5, 49, 166 Multi cycle 9ff NE German 48, 74ff, 126, 130, 162, 177, 183, 196, 306ff, 442, 443 North German 24ff, 48, 53ff, 73ff, 102ff, 119, 126ff, 152, 162ff, 181ff, 228ff, 233ff, 241, 267, 292, 300, 309ff, 347ff, 388, 389ff, 409, 436ff Northern Sudetic 173, 175 overfilled 219 Pannonian 6, 7, 50, 58 Paris 5, 205, 350 Pripyat-Dnieper-Donets-Donbas 6 pull-apart 5, 24, 109, 126, 159, 173, 181, 206, 383 Rift 10, 11, 34, 134, 162, 197, 259ff, 307 Saale 173ff Saar-Nahe 103, 173, 175, 184, 186 Sediment-filled 38 simple 3, 10, 12, 42 starved 5 sub- (halbes Buch) subsidence 8, 41, 48, 102, 147, 161, 187, 223, 244, 400 Thuringian Forest 173ff, 219 Underfilled 170, 219, 225 Basin and Range 45, 173 Bathymetric High 245 Bay of Kiel 77, 245 B-dot 405, 406 Bed-length balancing 99 Berriasian 167, 168, 207ff, 318, 415 Biodegradation 362, 413ff, 449ff Biogenic 171, 222, 396, 446 Biomarkers 139, 427 Bischofite 277, 284, 286, 353 Bitter Springs Fm. 261 Bittern salts 252, 254 Bitumen 138ff, 329, 330, 364, 401ff, 448ff Black oil 9, 362ff Black sea 8, 259, 413 Black shale(s) 28, 83, 186, 201ff, 438, 439 Bleaching 401ff Block Boundaries 233ff Block rotation 97, 206 Block-and-ash flow deposits 176 Bohemian Massif 6, 50, 104, 105, 120, 166ff, 203ff, 219 Boreal 159, 185, 209, 210, 230 Borehole break-out 120, 121 Bramsche Area 93, 224 Massif 93, 94, 456 Brandenburg 174ff, 185ff, 239, 319, 351 Brine chemistry 171, 252ff, 353ff evaporation 172, 521, 252ff, 258, 262ff, 279ff, 305, 323ff, 347ff, 370, 387, 390, 457 Brittle 46ff, 111, 163, 233, 278, 297, 347, 373, 380ff Brittle failure 50ff, 368,

Broad Fourteens Basin 129, 313 Buckling 5, 37, 48, 49, 66 Bunter 329, 353 Buntsandstein 28, 31, 75, 91, 109ff, 135, 148, 159ff, 181, 188ff, 219, 222, 228, 249, 250, 268, 269, 293ff, 312ff, 329, 338, 342, 365, 376, 379, 408, 410, 445 Burgundy Gate 192ff Burial diagenetic overprint 173ff, 367, 443 history 125, 152, 269, 326, 348, 408, 411, 417 Büsum 324ff Byerlee’s law 51ff, 61 C Calcite 117, 254, 277, 355, 356, 390ff, 451ff Calcite e-twins 117 Caledonian 19ff, 58, 68ff Caledonian Deformation Front (CDF) 19, 20, 58, 73ff Callovian 167, 201ff, 415, 416 Calorific values 140, 433 Campanian 109, 167, 210, 216ff, 301 Capillary entry pressure 330, 372ff, 421 Capillary seal 373ff Carbon content 138ff, 328, 359, 433, 439 dioxide 9, 12, 329, 347, 360ff, 401ff, 433ff preference index 139, 141 Carbonate platform 194, 205, 330 stringer 329, 330, 430, 431 Carboniferous 11, 17ff, 22ff, 64, 82ff, 93, 94, 103, 117ff, 125ff, 141ff, 157ff, 173ff, 182ff, 206, 223ff, 254ff, 311, 321, 327ff, 349, 363, 394, 395, 399ff, 415, 418, 433 Carnallite 254, 277, 284, 286 , 353, 354 Carnian 164, 194ff Cation 254, 349ff, 392, 441, 443 CELEBRATION 2000 19 Celestine 270 Cementation 112, 330, 348, 365ff, 399, 404ff, 429, 446, 447 Cenoman(ian) 112, 158ff, 220ff Cenozoic 26ff, 72ff, 105ff, 125, 131ff, 159, 169, 171, 173, 174, 187, 210, 221, 292, 301, 311, 315ff, 338, 339, 365, 390, 400, 434 Central Atlantic 125, 130, 159, 166, 235 Central European Basin System (CEBS) 5, 11, 17ff, 37, 42, 57, 62ff, 67ff, 97, 109, 125, 126, 141, 148, 157ff, 173, 174, 181, 233, 249, 270ff, 291, 304, 308, 323, 324, 337, 349, 365, 389, 412ff, 433ff, 456 Central Graben 11, 25, 26, 57, 60, 70, 71, 79ff, 104, 126, 129ff, 149, 164, 183, 192ff, 308ff, 387, 406, 407, 415, 428 Central Netherland Basin 159, 221ff Chalk 28, 131, 143, 161, 167ff, 213ff, 268, 356, 365, 390 Charge factor 10 Chloride 163, 278, 348, 355, 389, 390, 405, 409 Chlorinity 347. 348 Chlorite 409, 410 Chotts 271, 272 Chronostratigraphy 160, 199 Cimmerian 47, 65, 130, 160ff, 181, 195ff, 300 Classification of sedimentary basins 3, 32 Clastic(s) 28, 72, 77, 77, 83, 126ff, 157ff, 173ff, 182ff, 211, 215ff, 233, 271, 302, 308, 309, 356ff, 389ff, 407, 438, 447, 449 Clay 8, 9, 97, 136, 141, 171, 185ff, 218ff, 233, 235, 249, 270ff, 303, 328, 349, 357, 358, 368ff, 389ff, 415, 421, 433ff Clay diaper 8, 9

Subject Index Clay mineral 97, 141, 218, 357, 368, 387, 390, 391ff, 435ff Clay mineral – transformation 368 Clay smear 382, 385 Climate(ic) change 3, 251, 410 processes 256 zones 26, 258, 259 CO2-sequestration 410, 444 Coal(y) bearing 30, 141, 145, 153, 157, 167, 170, 327, 328, 418, 433ff seams 28, 85, 418, 435 Coalification 140, 220, 326 Coastal deserts 255ff Coastal Salinas 265 Coastlines 233 Coulomb-Navier 51 Common Midpoint (CMP) 69, 74, 75 Common Reflection Surface (CRS) 69, 74, 75, 82 Compaction 91, 98ff, 134, 137, 171, 187, 235, 279, 281, 297, 320, 331, 347, 348, 367ff, 386ff, 391, 399, 408, 409, 412, 427ff Compaction water 348 Compartment 9, 192, 211, 216, 218, 347, 387, 388 Complex basin 3ff, 34, 37, 42, 60, 66 Compound groups 362, 426, 427 Compression 5ff, 25, 26, 37, 47ff, 72ff, 109ff, 125, 132, 133, 161ff, 173, 210, 211ff, 258ff, 292ff, 307, 313ff, 331, 333, 368, 431, 445 Condensate 9, 359ff, 412, 422, 425, 428, 451, 455, 456 Condensed section 213, 215 Conglomerate Variscan basement 176 Conodont colour 139, 140 Continental margin 8, 41, 42, 136, 137, 259, 272, 305, 416, 421 crust 5, 37, 41, 49, 66, 136, 173, 265, 269 Continental encroachment 161 Continentality 256, 258 Contraction 5, 37, 99ff, 181, 387 Convective cell 332ff regime 340ff Convection forced 134, 334, 341 free 331ff gravitational 331 mixed 334, 341 thermal 134, 367, 387, 388 thermohaline 332ff, 386, 387 Continental margin 8, 41, 42, 136, 137, 259, 272, 305, 416, 421 crust 5, 37, 41, 49, 66, 136, 173, 265, 269 Cooling 4, 38ff, 101, 104, 128, 134, 142, 149ff, 166, 168, 173, 176, 176, 181, 187, 209, 224, 227, 256, 259, 334, 394ff, 432, 434 Cooling history 149 Coral reefs 28, 194, 201 Corrugations 117ff Cracking oil to gas 421, 430 secondary 420ff Creep steady state 52 transient 52 linear 59 Cretaceous Early 25, 28, 30, 47, 66, 72, 78, 104ff, 128ff, 159, 166, 169, 181, 206ff, 217ff, 254, 300ff, 308ff, 326, 401, 415, 444ff

Late 6, 7, 23ff, 48, 49, 58, 65, 66, 72ff, 93, 94, 103ff, 130ff, 158ff, 181ff, 212ff, 255, 268, 269, 296ff, 308ff, 365, 444, 446 palaeogeography 207 stratigraphy 208, 266 Cross-section balancing 99 Crude oil 31, 360ff, 380, 425, 448ff Crust(al) association 6, 13, 19 delamination 173 discontinuities 29, 322 domains 7, 19, 87 extension 41ff, 128, 130, 187, 206, 209, 266 heterogeneities 6, 25, 89, 95 Relaxation 215, 233 structure 3ff, 20, 30, 48, 67, 69ff, 321 thinning 4, 41, 42, 115, 149 Crystal plasticity 278, 280 Cryo-SEM 283, 285 Curvature dip 85ff Cycle coarsening-upward 161 depositional (DC) 157ff, 181ff east Tethys 166 evaporation 163, 185 evolutionary 161 North Sea 167 onlap-offlap 161 sedimentary 10, 157, 204, 435 transgression-regression 161 D Danian 161, 167, 219, 228 Danish Basin 24, 57, 60 ff., 87, 89, 104, 125 ff., 171, 193, 197 ff., 228, 231, 309 ff, 321 Danish lowlands 19 Davies 405, 409 Dead Sea 5,6 256 ff., 305, 390 Death Valley 5, 262 Debris flow 188, 218, 224 Debye-Hückel 405, 406 Decompaction 100, 297 Deconvolution Euler 70, 85, 89 De-coupling /decoupling 42, 43, 52 ff., 66, 109, 115, 212 ff., 266, 292, 297 ff., 308, 312, 329 Deep gas 12, 409, 422 Deep sea graben 5 Deep structure 19, 37, 81, 112 Deep time 249 ff. Deformation 4 ff., 19 ff., 38 ff., 69 ff. 97 ff., 125 ff., 181 ff., 211 ff., 235, 236, 277 ff., 307 ff., 328 ff., 368 ff., 394, 409 Contractional 105 Extensional 108, 210 Heterogeneous 107, 278 Mechanism 52, 281 ff. Polyphase 106, 119 Deglaciation 120, 236, 245, 357 Dehydration of clay minerals 368, 387, 390 DEKORP-BASIN’96 94 Delayed infill 43, 44, 128 Density 5, 8, 22, 26, 30, 38, 40, 47 ff., 69, 79, 84 ff., 100, 101, 135, 168, 254, 277 ff., 292, 301, 331 ff., 359 ff., 374 ff., 390, 423 ff., 458 model 89 ff.

509

510

Subject Index Depocentre /Depo-centres 17, 24 ff., 49,60, 64, 66, 94, 104, 128, 129, 131 ff., 162, 163, 169, 181 ff., 206, 211 ff., 249, 293 ff., 310 ff., 403 Deposition(al) 8. 9, 13, 28, 73, 77, 78, 97, 122, 126, 128, 130 ff., 142, 149, 151, 157 ff, 179, 181 ff, 212 ff, 249 ff, 292 ff, 307 ff, 330, 348, 367, 369, 389, 396, 399, 408, 409, 412 ff, 426, 438, 441, 443, 447 Cycle 161 ff, 181, 184, 195, 204 Environment 8, 97, 169, 170, 182, 186, 189, 194, 201, 207 post-tectonic 8, 9, 307, 308, 321, 447 pre-tectonic 168 syn-tectonic 122, 130, 164, 167, 187, 269, 293, 303 ff system 162, 169, 171, 172, 269 Desert (s) 188, 255 ff, 272 ff, 434 Desert belts 256 Desiccation 270, 399 Detachment 6, 48, 56, 59, 105 ff, 212, 233 Devonian 6, 47, 92, 139, 140, 260, 261, 264, 311, 408, 418, 448, 451 late 6 Diagenetic, Diagenesis 8, 9, 12, 13, 28, 99, 133, 176, 178, 330, 335, 348 ff, 355, 360, 367 ff, 387, 388, 390 ff, 420, 446 ff process(es) 8, 12, 28, 99, 349, 351, 368, 369, 387, 388, 391 ff, system Diapir (s) 8, 9, 13, 30, 38, 53, 59, 71, 73, 90, 94, 100 ff, 131, 146, 170 ff, 187, 190, 198, 210, 214, 219, 220, 230, 233 ff, 249, 268, 269, 273, 274, 285 ff, 291 ff, 308 ff, 323 ff, 335, 339 ff, 358, 388, 449 Differential loading downbuilding 292ff passive diapirism 292ff stress 293ff Diffusion 279ff, 332ff, 347, 367ff, 390, 407, 436, 450ff Digital Elevation Mode 233ffl Dilatancy 278ff Dilatation 111, 113, 278 Dinarides 50 Dip 5, 34, 45, 46, 68ff, 98ff, 221, 226, 268, 275, 294ff, 316, 385 Dip slip 5, 108 Discontinuity 44, 59, 269, 370 Discrepant zone 45, Dislocation 7, 59, 277ff Displacement Distribution 117 Dissolution 100, 170, 192, 214, 244, 255, 258, 261, 262, 268, 273 ff, 277, 280, 287, 290, 297, 301 ff, 322, 336, 338, 339, 347 ff, 368, 369, 384, 387, 388, 389 ff, 399 ff, 431, 446, 447, 452, 454 Dniepr-Donets 48, 260, 261, 264 DOBRE-flection 6, 47, 48 Dolomite 18, 133, 135, 164, 171, 192 ff, 270, 277, 369, 371, 390, 400, 410, 451, 452, 455 Dolomitization 390, 447 Domal salt 287 Donbas 6, 7 47, 48 Dorn’s law 61 Down-building 8 Dowsing–South Hewett Fault Zone (DSHFZ) 20, 86, 87 Drainage Pattern 194, 234, 243 Drilling induced tensile fractures (DITF) 121 Ductile 29, 43, 44, 51 ff, 61, 78, 11ff, 129, 277, 328, 372, 380, 381, 409 Dutch lowlands 19 E E Brandenburg Sub-Province 174, 177 Early Cimmerian 160ff, 181ff, 300

Earthquake focal mechanisms 119 Earthquake(s) 43, 67, 70, 81, 86, 119ff, 193, 236, 241, 387 East African Graben System 4 East Avalonia 19, 20 East Carpathian Gate 187ff East European Craton 19ff, 60, 61, 87, 125, 143 East-European Platform 259 Eichsfeld-Altmark swell 104, 111, 163, 164, 197, 199, 304 Ekofisk 34 Elastic beam 37, 49 plate 48, 50 response 235 stiffness 50 Elasto-plastic 53 Elbe Fault System (EFS) 20ff, 57ff, 86, 87, 109, 126, 143, 308ff Elbe-Lineament 234 Elbe-Odra Line (EOL) 19ff Element distribution 355 Elsterian Glaciation 234, 238, 432 Enhanced oil/gas recovery 410 Entrapment 10, 31, 392ff, 401, 457, 458 Eocene 4, 122, 124, 143, 159, 161, 167, 169, 230, 231, 274, 289, 356 Epeiric sea 169, 218, 255, 270 Equilibrium system 405 Erosion 5ff, 26ff, 38, 44, 45, 72, 77, 78, 100ff, 130ff, 143ff, 158, 162, 165ff, 183ff, 212ff, 233, 237, 243, 258, 262, 292, 299ff, 307ff, 358, 367, 368, 432, 446 EUGENO-S 19, 61, 79, 80 Eurasia 25, 28, 50, 56, 125, 130 Eurasian margin 130 Eurasian plate 50, 56 European Cenozoic Rift System 123 European Geotraverse (EGT) 19, 79ff Euxinic 163ff, 186, 201 Evaporation 17, 163, 171, 185ff, 251, 304, 349ff, 389, 392, 399, 406, 409, 443 Evaporite cycle (s) 28, 188, 192 Evaporite(s) 28, 31, 70, 77, 126ff, 126ff, 159, 163ff, 182ff, 222, 228, 241, 249ff, 277ff, 293, 302ff, 307, 308, 329, 330, 349ff, 372, 389ff, 415, 418, 430, 447, 449 Evaporitic sediments 249 Exhumation 219ff, 401 Extension (al) post orogenic 173 F Facies complex 8 Failed rift 37 Failure7, 48ff, 121, 218, 233ff, 278, 368, 380, 382, 421, 431 Fault array 103ff blocks 201, 233 crustal scale 46 cut-off 114 dip 112, 115 extensional 45, 181, 183, 190, 266, 269, 303, 304 gouge 381ff graben 109 intracratonic 119 lystric 7 mesoscale 118 normal 64, 103ff, 128ff, 159, 206, 211, 220, 239, 243, 266, 293ff, 310ff, 385

Subject Index oblique 5, 111, 300 offset 103 plane diagram 381, 383 population 117ff reactivation 106, 124, 211 reverse 26, 105ff, 167, 190, 220, 301, 305, 315, 319, 444 seal 372ff set 112ff, 293 slip direction 117, 118 strike-slip 64, 99, 103ff, 159, 162, 170, 206, 292 synsedimentary 112, 189, 210 system 5, 20, 22, 24, 26, 48, 57ff, 97ff, 126, 132, 187, 210, 231, 235, 304, 308ff, 381, 398 throw 112, 382, 383 thrust 113 tip 383 transtensional 24, 104, 121, 206 zone 19, 20, 59, 75, 86, 87, 102ff, 132, 133, 143, 226, 233, 245, 309, 347, 372, 377, 381ff, 391, 997, 404, 429 Fault-bounded Blocks 100, 233 Faulting, active 43, 120 Fault-propagation fold 112, 214, 228 Fault-slip analysis 117, 118 Fault-Surface-Penetration-Point 233 Feldspar 228, 229, 349ff, 388, 390ff, 435ff Fennoscandia 72, 159, 161, 169, 194ff, 234, 236 Fennoscandian Ice Sheet 234 Field size distribution (FSD) 9, 31ff Fining-upward 162, 163, 184ff Finite-element 50, 61 Finland 56, 236 First order approximation 42, 51, 126 Fission track(s) 112, 141, 142, 149ff, 165ff, 186, 197, 223ff, 391, 397, 398 Fissure mineral 392ff Fjerritslev Trough (FT) 25, 26 Flechtingen High 102ff, 221, 318 Flechtingen-Altmark Sub-Province 174, 177 Flechtingen-Roßlau Block 173ff Flexure(al) bulge 48, 113, 212 model 48ff rigidity 49 Flinders Ranges 261, 271 Flow law 98, 277ff, 384 Fluid chemistry 285, 397, 406 compartments 9, 388 composition 30, 391, 399, 406, 422ff density 331ff, 382, 424 entrapment 392, 457, 458 evolution 349, 389ff flow 112, 134, 278, 286, 287, 323ff, 347, 349, 367ff, 381ff, 390, 391, 398, 399, 407, 428 inclusion (s) 141, 152, 253, 254, 284ff, 348, 390ff, 429, 439, 455ff inventory 3, 9, 34 migration 12, 330, 349, 391, 401 mixing 406 reconstruction 391, 399 rock-interaction 336, 347, 389ff, 435ff Fluorescence 139, 140, 392 Fluvial erosion 233, 237 system 8, 12, 197

Fold(ing) 5, 7, 97ff, 162, 168, 179, 187, 188, 212ff, 260ff, 285ff, 292, 293, 299ff, 312ff, 418 Fold-fault-propagation 112, 214, 228 Forearc 49 Forebulge 120, 212, 213, 231, 233, 238 Foreland 6, 22, 37, 94, 97ff, 134, 162, 166, 179, 181, 211ff, 238, 258, 260ff, 307, 425 Foreland deformation 214 Formation water 134, 347ff, 388, 390ff, 452ff Fossil forest 28 soil 28 Fractionation 356, 393ff, 427, 435ff Fracture porosity 173, 372 Fractures 7ff, 51, 72, 98, 117ff, 173, 175, 187, 206, 245, 286, 287, 328, 330, 344, 347ff, 367ff, 391, 398ff, 446, 453 Fracturing 112, 121, 286, 287, 328, 330, 370ff Franconian tectonic movements 179 Friction 51, 56, 117, 380ff Friesland Platform 143 Front limb 109, 112 G Gabelsflach 243ff Galicia 41 Gamma-irradiation 285, 287 Gardelegen Fault 75, 109ff Gas biogenic 446, 451 condensate 363, 422, 428, 451, 456 deep 12, 409, 422 filling 329 generation 133, 137, 153, 324ff, 360ff, 400, 414ff, 439ff hydrate 419, 422, 430, 431 inorganic 9, 347, 364 microbial 364, 418, 421, 431 migration 329, 347, 400, 431, 436, 458 natural 9, 125, 347, 348, 360ff, 408, 410, 412, 433ff reservoir 28, 30, 145, 328, 329, 365, 399, 400, 408, 410, 434ff shallow 418, 419 thermal 360, 363, 418, 440 thermogenic 364, 419, 457 tight 409, 422 Gas field (gasfield) 18, 31ff, 327, 386, 387, 400, 410, 419, 445ff Gas prospect 32 Gaussian curvature 118 Generation Gas 133, 137, 153, 324ff, 360ff, 400, 414ff, 439ff Hydrocarbon 31, 165, 323, 331, 411, 419ff, 432 Oil 31, 327, 360ff, 404, 414ff Geochemical reactor 9, 19 Geofluid 347ff Geophysics (al) investigation 67 methods 67, 72, 95, 125 Geo-reactor 9, 125, 347 Geotherm(al) energy 133, 347, 410 field 56, 331ff gradient 46, 133, 134, 151, 227, 342, 344, 425 German Midlands 19 Glacial landforms 233, 237 loading 119, 431, 432 Glaciation (s) 28, 29, 120, 188, 233ff, 357, 358, 430ff

511

512

Subject Index Glaciofluvial 233, 237 Glaciolacustrine Global Positioning System (GPS) 121 Global sea level 5, 8, 26, 130, 161, 162, 169, 186ff, 206, 210, 223 Glückstadt Graben (GG) 25, 26, 57, 60, 64, 67, 72ff, 102, 122, 124, 126ff, 148ff, 164ff, 184, 190, 197ff, 230, 240ff, 300, 305, 308ff Gondwana 94, 188, 259, 261 GPS 56ff, 120, 121 Graben shoulder 83, 159, 170, 206 Grain boundary 277ff, 330 Grain boundary migration 279ff Gravitation 83, 171, 218, 289, 307ff, 331ff, 367 Gravity field 84ff, 338 inversion 292 residual 87ff Great Artesian Basin 271, 272 Greenhouse 159, 218, 251ff, 410, 422, 444 Greenland-Norwegian Sea 128 Groundwater chemistry 347 Groundwater flow 264, 323, 330ff, 341, 343, 347, 358, 374 Great Glen fault 5 Greenland 28, 64, 128, 185ff, 259 Growth fault 293 Growth structure 215 Gulf coast region 260, 264, 336, 388, 454 Gulf of Mexico 254ff, 291, 293, 305, 323, 324, 390 Gypsum 28, 164, 169, 196, 206, 241, 245, 252ff, 277, 389, 390, 401, 405, 406, 452 H Hadley Cells 255ff Haldensleben Fault 75, 109 Half graben 112ff, 175, 184, 205 Half ramp basin 212 Halite 28, 83, 163ff, 249ff, 277ff, 301, 305, 328ff, 389ff, 447 Halite stratigraphy 250 Halokinesis 75, 250, 255, 266ff, 292 Halokinetic styles 268 Halotectonic 160ff, 186, 194 Haloturbation 270 Hannover 17, 18, 93, 121, 163, 176, 182ff, 444 Hardegsen 160ff, 181, 189, 190 Hardground 215ff Harz 17, 19, 48, 75, 102ff, 167, 170, 204, 219ff, 238, 395 Harz Boundary Fault 48, 75, 114 Harz Mountains 17, 19, 109ff, 167, 170, 204, 219ff, 238, 395 Havel 135, 162, 182ff, 233, 337 Heat conduction 39, 41, 61, 133, 323, 455 convection 133 flow 5, 38ff, 125, 133ff, 184, 233, 241, 334, 343, 367, 387, 413, 456 loss 41, 42 production 39, 61, 125, 135, 136 production-radiogenic 39, 61, 125, 135 radiogenic Heating problem 38 Heavy minerals 168, 219 Hematite 111, 443 Hennstedt 239 Hercynian basin 110ff, 129ff, 170, 221ff, 313

Hessian Depression 186ff, 231 Heterogeneity(ies) 6, 7, 25, 50, 59ff, 68, 89, 95, 109, 181, 335, 368, 414 Hiatus 167, 169, 180, 215, 218, 225 High velocity body 6, 7 Holocene 235, 236, 244, 245, 255ff, 431 Holy Cross Mountains 19, 102, 104, 309, 315 Hopper grains 287 Hormuz Salt 261, 275 Horn Graben (HG) 25, 28, 57ff, 70, 73, 86ff, 104, 126ff, 162ff, 176, 187, 197, 206, 249, 308ff, 321 Horse latitude deserts 259 HPHT 428 Husum 239 Hydraulic fracturing (hydrofracturing) 121, 287, 328, 330, 374, 375 Hydrocarbon(s) exploration 18, 142, 173, 291, 348, 399, 412, 444 generation 31, 165, 323, 331, 411, 419ff, 432 maturation 142 migration 386, 392, 419 system 10, 32 Hydrodynamic trapping 374 Hydrogen sulphide (sulfide) 9, 12. 363, 364, 401, 433, 447 Hydrographic isolation 264, 265, 269, 302 Hydrological regime 3 Hydrothermal 111, 113, 134, 138, 146, 149, 173, 178, 251, 357, 380, 391, 394ff, 434, 443, 449, 457, 458 Hydrothermal overprint 173, 178, 394, 443 I Iapetus Ocean 85 Ibbenbüren High 219 Ice Marginal Valley(s) 29, 233 Ice sheet 29, 120, 159, 169, 233ff, 251, 257, 285 Ice sheet- crust-interaction 238 Ice thickness 235 Ice wedges 233 Icehouse 251, 254, 257, 259, 276 Igneous dykes 147 Ignimbrite non-welded 176 sheet 174ff, welded 24 Illawarra reversal 182, 183 Illite 111, 368, 388, 391, 397ff, 435ff Illitisation 387, 391, 398, 407 Indenter 105 Inherited 6ff, 25, 29, 34, 59ff, 132, 239, 243, 356, 426 Intra-plate deformation 48, 56, 59 stress 47ff, 130 orces 50 Interfacial tension Interferometric Synthetic Aperture Radar (InSAR) 121 Intrusion 38, 90ff, 146, 149, 176, 177, 293, 456 Inversion 5ff, 25ff, 37, 42, 48, 49, 60, 66ff, 103ff, 125ff, 161, 166ff, 181, 184, 201, 206, 210, 211ff, 268, 287, 292, 301, 306, 308ff, 338, 343, 400, 401, 409, 418, 444ff, 458 Inversion structures 103, 123, 132, 217ff, 313, 315 Iron oxide 399, 401 Ironstones 217ff Isochore 77, 392ff, 457 Isostasy 100, 101, 203, 236, 245, 297 Isostatic Adjustment 56, 233, 245 Isostatic Equilibrium 38, 99, 101, 235, 236, 321

Subject Index Isostatic Rebound 59, 233, 235, 238, 320, 431 Isotope 142, 157, 189, 355ff, 389ff, 421, 427, 434ff Isotopic composition carbon 360, 364, 395, 427, 446, 447, 453, 457 oxygen 157, 357, 358, 394ff, 453 strontium 357, 398 sulphur 358, 452 J Jet stream 258 Jordan rift 5 Jurassic Early 31, 64, 65, 72, 77ff, 111, 130, 159, 166, 196ff, 202ff, 294, 308ff, 321, 416 Late 25, 28, 47, 66, 92, 104, 108, 122, 130, 131, 143ff, 158, 159, 166ff, 199ff, 220ff, 254, 300, 308ff, 365, 401, 415, 418, 444 Middle 65, 72, 77, 122, 130, 143ff, 159, 166, 167, 181, 199ff, 302, 306, 325, 457 K K/Ar dating 111, 176, 443 Kalahari 255, 256 Kaolinite 388, 390, 391, 400, 401, 404, 435 K-Ar age 162 Kerogen classification 361, 413, 414 transformation 361 type 360, 364, 415 Ketone(s) 157 Keuper 28, 78, 108ff, 122, 135, 148, 164, 166, 170, 171, 181, 191ff, 249, 250, 268ff, 294ff, 310ff, 325, 338, 410 K-feldspar 349, 357, 388, 390, 391, 397, 398, 406, 407 Kimmeridgian 167, 203ff, 415, 416 Kinematic detachment 56 Kinematic(s) 4, 37, 38, 50ff, 75, 97ff, 108, 111, 117, 118, 124, 293, 297ff Kinetic (s) 125, 238, 245, 261ff, 278, 292, 293, 304, 308, 324, 390, 405ff, 411ff, 420ff, 436ff Klemme’s approach 31 Klüpfel 161, 167 Kuh-e-Namak-Qom 273ff Kupferschiefer 83, 163, 185ff, 408 Kyffhäuser 111, 219 L Labrador 42 Laccoliths 176 Lacustrine halites 258 Ladinian 164, 191ff Lake Eyre basin 271 Lake Plön 240ff Landscape 9, 29, 233ff, 269, 275ff, 305 Landscape evolution 119, 120, 179 Late Cimmerian 160, 167, 181, 207ff, 300 Lateral variation 6, 13, 34, 42, 49ff, 56ff, 292, 297, 321 Laurentia 79ff, 85, 94 Laurussia 19, 259, 261 Lausitz high 219 Lava flow andesitic 178 basaltic 178 rhyolitic 174 Lava dome 174, 176ff, 180 Lavadome-related pyroclastic deposits 176 Layering 46, 53, 303, 369

Lettenkeuper 28 Levante rift system 5 Liassic Posidonia Shale 153, 327, 365 Lignite 18, 133, 138, 141, 169, 230, 233, 241 Ligurian 167 Limestone 18, 28, 122, 159, 163, 166ff, 171, 187, 191ff, 203, 205, 219, 223ff, 270, 275, 349, 355ff, 391 Lineament(s) 81, 85, 103, 109, 126, 190, 197ff, 233ff, 237ff, 313, 318 Liquefaction 236 Lithology 24, 100, 121, 133ff, 168, 179, 250, 278, 325, 355, 372 Lithosphere(ic) barrier 38ff, 434 continental 5ff, 13, 19, 50ff, 119, 125, 173, 234ff cooling 79ff, 128, 166 extension 38ff, 50ff, 79ff, 98ff, 119, 212ff oceanic 6, 37, 234ff Lithostatic pressure 287 Loch Ness 5 London-Brabant Massif 141, 196, 267 Lower Saxony 129ff, 180, 196, 237, 312, 327 Lower Saxony Basin (LSB) 17, 25, 28, 34, 75, 104ff, 107, 122, 129, 143ff, 165ff, 200ff, 205, 217ff, 312, 329, 415ff, 434, 444, 457 Lüneburger Heide 17 M Magma(tism) 43, 64, 94, 104, 126ff, 162, 173ff, 183, 201, 206, 331, 394, 435ff Magmagenesis late Paleozoic volcanic rocks 173ff Magmatic underplating 94 Magnetic field 22, 82ff Magnetite 399, 402 Magnetotelluric 18, 75ff Mantle lithosphere 38ff, 173 Mantle plume 7, 173 Marginal trough primary 212 secondary 168, 211, 212, 228 Marine Seismic Surveys 245 Mass balance 169, 325, 329, 406, 450, 452 Mass flow 171, 218ff Maturation 30, 137ff, 323ff, 360, 364, 391ff, 411ff, 439ff Maturity 31, 93, 94, 125ff, 165, 219, 224, 323ff, 364, 408, 413ff, 438ff Maturity gradients 146 Maturity parameters 137ff McKenzie’s model 41, 42 Mecklenburg Bay 78 Mecklenburg-Vorpommern 120, 174ff, 239 Mecklenburg-Vorpommern Sub-Province 174, 177 Mediterranean 7, 49, 56, 187, 230, 260ff Megahalite basins 266 Meltwater erosion 233, 243 Meltwater sediments 233 Messinian evaporates 261 Metamorphic process 13, 24, 85 Metamorphism 5, 24, 139, 278, 287, 348, 349, 435ff Meteoric water 9, 134, 192, 251, 264, 276, 348ff, 386, 388, 389ff, 426, 434, 453 Methane bacterial 360 thermogenic 364 Methane generation 327, 360, 364, 421, 436, 439, 446 Method

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514

Subject Index backstripping 85 backward 86 forward 86 Methyl-Phenantrene Index 139, 140, 141 Microbial degradation 359, 421 Microcrack (ing) 278ff, 330 Microthermometry 391ff Mid-Atlantic Ridge 53, 56, 237 Middle Cimmerian 160 Mid North Sea High 130, 159, 186, 193, 267 Mid Polish Swell 131, 230, 231, 308ff Migration pathway 356, 359 Milankovitch 161ff, 185ff Milankovitch cycles 161, 185, 194 Mineral alteration 348 Mineral deposit 408 Mineral inventory 9 Mineralisation 352, 392ff, 457 Minimum principle stress 374 Miocene 7, 26, 124, 150, 169, 232, 243, 254ff, 433 Miocene salt 265ff Mississippi-Valley-type 408, 447, 452 Mittelplate 324ff, 418, 431, 432 Model(ling) 2D 98 3D 84, 92, 107ff, 325, 327, 418 dynamic 54ff, 2236 flexural 48ff kinematic 37, 38, 53, 54, 98 McKenzie’s 41, 42 non-uniform stretching 42ff numerical 37, 55, 98, 137, 153, 165, 285, 324ff, 383, 391, 411 pure shear 44ff simple shear 44ff spherical 56ff thermal 37, 66, 140, 400 thin-plate 55, 59 thin-sheet 55ff Moho topography 49, 74, 89 Mohr circle 383 Mohr-Coulomb yield criterion 117 Mohr diagram 382 Molecular geochemical parameters 138, 141 MONA LISA 19, 61, 71ff Moraines 233ff Moray Firth Basin 65, 200 Multi parameter stacking 68, 69 Multiple Inverse Method 119 Münsterland Basin 217ff Müritz (group) 135, 162, 179, 182 Muschelkalk 28, 91, 109, 118, 135, 159ff, 188ff, 249ff. 294ff, 309ff, 340, 342 N Namakier (s) extension 249ff middle Keuper 302, 303 terrestrial 305 Natural gas 9, 125, 347ff, 408, 410, 412, 433ff NE Brandenburg - Wolsztyn High 180 Necking depth 49 Neotectonic period 118 Neotectonic(s) 77, 118, 321

Netherlands 32, 102, 104, 132ff, 163ff, 181ff, 233, 238, 285, 287, 313, 399, 401, 418, 449, 457 Neuhof 287, 288 Niger delta 34 Nitrogen 9, 12, 30, 153, 362ff, 401, 403, 433ff Nitrogen generation 439ff Non-deposition 78, 158, 167, 213, 215 Non-linear process 37, 335, 412 North Atlantic 125, 159, 161, 168, 184ff, 254, 261, 267, 292 North East German Basin (NEGB) 48, 49, 165, 204, 323, 330, 337 North German Basin (NGB) 24ff, 48ff, 73ff, 102ff, 125ff, 181ff, 228ff, 235, 236, 241, 267, 300, 309ff, 325, 347, 348, 350, 389ff, 436ff North German lowlands 18 North Sea 10, 11, 19ff, 45ff, 68ff, 104ff, 125ff, 159ff, 174ff, 181ff, 231, 233, 266ff, 292ff, 307ff, 325, 365, 388, 400ff, 415ff, 433 North Sea Basin 10, 25, 49, 72, 161, 169, 181, 189, 192, 312 North Sea Dome 130, 203, 204, 206 North Sea High 130, 159, 186, 193, 267 North Sea rift 65, 66, 104, 130, 166, 206, 209, 210, 300, 305 Northern Permian Basin 17, 64, 126ff, 173, 183, 206 Norway 19, 28, 89, 105, 181, 235, 431 Norwegian Sea 128, 130 Norwegian-Danish Basin (NDB) 60, 87, 89, 104, 125ff, 200, 309ff Nova Scotia 41, 42 Numerical Dynamic Analysis (NDA) 119 Numerical modeling 137, 165, 324, 383, 391, 411 O Oberrhein Graben (Upper Rhine Valley) 4, 6, 136 Ocean chemistry 254 Ocean gyres 259 Oceanic Anoxic Event (OAE) 166, 167, 201 Oil field 18, 32, 325, 347, 348, 365, 418, 431, 453 flow 330 formation 360 generation 31, 327, 360ff, 404, 414ff maturation 323 migration 403ff prospect 32, 417, 418, 429 reservoir 69, 328, 364, 408, 432, 453 Officer Basin 260ff Offlap 13, 48, 161, 172 Oil-source rock correlation 138 Oligocene 66, 119, 122, 169, 230, 231, 289, 433 Onlap 8, 13, 44, 48, 70, 109, 161, 187, 266, 300ff Oolith 163ff, 192, 205 Ordovician 19, 21, 83, 94, 261, 330 Organic acid 400ff, 449 Organic geochemistry 12, 414 Organic matter maturation 30, 137, 140, 323ff, 422, 439 Organic-inorganic interaction 401, 403 Orogenic belt 53, 261, 265 Orogenie (s) 19 Oslo Graben 103, 173, 187, 197 Osnabrücker Bergland 17 Outcrop data 98, 189 Overburden 10, 29, 51, 53, 214, 227, 275, 285, 292ff, 307ff, 340, 368, 383, 386, 411, 430 Overcoring 121 Overpressure 347, 374ff, 427ff Oxfordian 158ff, 181ff, 415 Oxygen 157, 355ff, 394ff, 413, 414, 450, 453

Subject Index P Palaeo bathymetry 19, 84, 89 climate 157 environment 157 geography 128, 157, 168, 182ff, 249, 270 latitude 157, 199, 249, 257, 259 piezometry 289 stress 97ff topography 98, 101 water depth 98ff Palaeocene 104ff, 143, 158ff, 229, 230 Palaeogene 26, 72, 122, 124, 131, 158, 168, 228, 230, 301, 306, 311, 316ff, 343, 408 Palaeozoic 17ff, 50,61, 64, 72ff, 107ff, 153, 159, 173, 176, 190, 226ff, 254ff, 301, 329, 389ff, 412, 430, 438ff Pangea (Pangaea) 22ff, 70, 125, 161, 188, 254ff Pannonian basin 6, 7, 50, 58 Passive continental margin 272 Passive diapirism 292ff P-B-T-Method 118 Pelagic 167ff, 213ff, 438 Peneplain 167, 223 Penetration Point 243, 244 Penetrative strain 97, 106 Perennial lake 287 Peripheral sink 305 Permafrost 233, 241, 419, 431, 432 Permeability 9, 30, 31, 277ff, 328ff, 347, 367ff, 399ff, 422ff Permian basin 17ff, 42ff, 68ff, 108, 126ff, 141ff, 159, 173ff, 182ff, 291 early/lower 17ff,, 43, 64, 126ff, 159ff, 181, 187, 259, 311, 319, 352, 353 magmatism 43, 64, 104, 149, 152, 173ff upper/late 112, 126, 170, 173, 311, 319, 320, 357, 358, 401 Permo-Carboniferous 83, 126, 127, 161, 175, 257, 403, 433ff Petrography 168, 392 Petroleum classification 10, 359ff generation 10, 137ff, 327, 359ff, 404, 414, 416, 432 fluid (s) 30, 31, 347, 359ff, 401, 404, 422ff Petroleum system modeling 411ff Petrology 53, 81, 168, 456 Phanerozoic 82, 161ff, 253ff, 396 Phase behaviour 362, 423 Piesberg 224 Piggy back basin 212, 213, 265 Pitzer equations 405, 406 Plagioclase 349, 390 Plane-strain 99 Plastic 53ff, 233, 278, 280, 328, 368, 380 Plastic failure 56 Plasticity 37, 278, 280, 383 Plate convergence 235 Plate tectonic 3ff, 26, 50, 53, 64, 77, 120, 121, 236, 255ff, 307, 321 Playa (s) 162ff, 179, 189ff, 256ff, 301, 400 Pleistocene 59, 119, 120, 159, 233, 241, 264ff, 357, 358, 431 Plume (s) 7, 173, 230, 332ff Poland 19, 28, 32, 63, 79, 105, 126, 159ff, 173ff, 181ff, 229, 233ff, 249, 254, 399, 408, 436 Polish basin 87, 102ff, 125ff, 163, 200ff, 230, 309ff, 433 Polish Trough (PT) 17ff, 57ff, 89, 115, 119, 130, 164, 165, 193ff, 228, 249, 267, 309ff POLONAISE’97 19, 79

Pompeckj Block 28, 104ff, 130ff, 165, 166, 209, 210, 312, 318, 433, 444 Pop up structure 6, 48 Pore fluid 9, 330, 348, 368ff, 389ff Porosity 8, 30, 32, 68, 100, 133ff, 173, 241, 277, 278, 328ff, 367ff, 400ff, 411ff Posidonia Shale 31, 32, 153, 166, 201, 203, 325, 327, 365, 415ff Post-glacial rebound 120 Postglacial uplift 119 Post-rift 39ff, 74, 128, 266 Post-rift subsidence 42, 44, 128 Potential surface 3 Power law 51ff, 282, 285 PQ2 76ff Precambrian 19, 22, 80, 92, 125 Precipitation 99, 137, 163, 171, 185ff, 220, 251ff, 277ff, 332, 347, 349, 353ff, 387, 388, 289ff, 446ff Pre-Permian surface 125, 126 Present-day stress 59, 77, 98, 119 Pressure solution 51, 99, 277ff, 368, 409, 447 Prignitz 207, 208, 221 Prignitz-Lausitz High 221 Principal stress maximum 119 minor 112, 119 Pripyat-Dnieper-Donets-Donbas 6 Progressive unconformities 213ff Proterozoic Europe 22, 82, 169, 262 Provenance analysis 168, 228 Pull-apart 24, 173, 383 Pull-apart basin(s) 5, 109, 126, 159, 181, 206 Pure-shear 44ff Push Moraine 233ff Push-up structure 170 PVT 3, 362, 375, 376, 423, 431 Pyrite 360, 401, 452 Pyrobitumen 360, 364 Pyroclastic deposits 171, 174, 176 Q Quartz 52, 117, 135, 205, 218ff, 357, 368, 390ff, 429, 435 Quaternary deserts 257, 272 glaciation 28, 237, 245 sediments 109, 233, 355 Qom Basin 273 R Radiation 82, 133, 255, 285, 287, 392 Radiogenic heat 61, 125, 135 Raft tectonics 292, 293 Rain shadow 255 Rank 9, 138ff, 328, 372, 417, 419, 427, 439, 440 Rare earth elements (REE) 338, 355, 356 Rayleigh 292, 333ff, 427, 442, 457 Reactivation 6, 43, 59, 60, 95, 104ff, 132, 133, 190, 211, 239, 306 Receiver functions 68ff Recrystallisation 277ff, 446 Red bed 191, 399ff Red Sea 5, 258ff, 380 Red Sea Rift 4 Redox-reaction 447, 451 Reducing 401, 447ff

515

516

Subject Index Regression (ive) 8, 159, 160ff, 185ff Relay ramp 383 Releasing bend 383, 385 Rerik 78, 94 Reservoir properties 408, 409 quality 371, 399, 409, 447 rock(s) 10, 28, 93, 138, 139, 2006, 328, 365, 367, 389ff, 411ff, 446, 447, 457 seal 382 trap 10ff, 31, 214, 268, 278, 325, 329, 348, 364, 372ff, 411ff Restoration 98ff, 125, 297ff, 312, 325 Retention 372, 375, 417, 428, 430, 446 Retro-deformation balancing 292, 305 modelling 297, 299 sequential 296, 297 Reverse fault (s) (ing) 26, 105ff, 167, 190, 220, 301, 305, 315, 319, 444 Rheic Ocean 20 Rheinsberg Trough (RT) 25, 26, 86ff, 102, 122, 126, 130, 309ff Rhenish Massif 105, 169, 170, 186ff, 223, 225, 267, 444 Rheology(ical) contrasts 6, 34, 63 Rhyolite 24 Ridge push force 50, 56, 59, 119 Ridge push 50, 56, 59, 119, 237 Rift (ing) event 41, 42, 184 simple 37 stuck 5, 13 zone 5, 41, 46, 197, 198 Rifting 5, 6, 24ff, 39ff, 80, 94, 123, 125ff, 159ff, 181ff, 249ff, 293ff Rim syncline (s) 8, 28, 108, 130, 198, 200ff, 273, 292ff, 312ff, 325ff Ringkøbing-Fyn High 19ff, 57ff, 73ff, 102ff, 126, 186ff, 308 River system 28, 29, 189 River (s) 8, 28, 29, 140, 171, 189, 218, 219, 233ff, 260, 264, 337, 338, 386, 413, 419, 448, 454 Rock mechanics 53, 380 Rock-Eval – hydrogen index 139ff, 360, 414 Rock-Eval – mineral matrix effects 141 Rock-Eval – production index 139ff Rock-Eval – Tmax 139ff, 414 Rocksalt 282ff Rodinia 261 Roenne Graben 5 Roer Valley 132, 143, 200, 204 Röt Basin 254 Röt brines 254 Röt halite 254, 269 Rotliegend 11, 17ff, 83, 91, 103ff, 126ff, 158ff, 173ff, 181ff, 234, 293ff, 307ff, 325ff, 353ff, 390ff, 433ff S Saalian Glaciation 245 Sabkha 163, 164, 183ff Sahara 255, 256 Saline giants 251, 261, 265 Salinity 214, 218, 253ff, 332ff, 347ff, 387, 388, 389ff, 413 Salt allochthons 305 anticline 109ff, 299 basin 187, 254ff, 291ff, 307ff, 328 brines 9, 171, 172, 251ff, 279ff, 305, 323ff, 347ff, 370, 387, 390ff, 443, 457

buoyancy 292 canopies 293, 305 crests 305, 339 décollement 269 diapir (ism) 9, 30, 59, 71, 94, 103, 107, 114, 131, 146, 187, 198, 210, 230ff, 238ff, 285, 291ff, 308, 318, 323ff, 352, 358 dissolution 274, 301ff, 389 dome 30, 77, 89, 171, 241, 288ff, 314, 323ff, 348ff, 387, 390, 447, 453 dynamics 77, 239, 247, 291, 304 extrusion 290, 302ff, 322 flanges 302, 304 flow 29, 267ff, 277, 285, 292ff, 317ff geometries 293, 297 glaciers 267, 268, 273, 285ff jags 300, 305 kinematics 268, 293, 297 lake 162, 163, 271, 390 migration 59, 109, 113, 214, 215, 293, 338 mobilisation 125ff, 307ff mother brines 251 movement 8, 28ff, 53, 71, 78, 91, 101, 105, 114, 123, 131, 172, 190, 200, 201, 268, 292ff, 308ff, 324 namakier 249ff, 302ff pillow 53, 78, 94, 106, 109, 170, 198, 214, 234ff, 268, 292ff, 308ff, 324ff re-sedimentation 300, 305 restoration 101 rise 130, 132, 268, 290ff, 310ff sediment interaction 301 sequence stratigraphy 304 sheet 292, 332 springs 337 stocks 293, 318 structure 8, 26ff, 53, 61, 74ff, 100ff, 129, 130, 239ff, 267ff, 291ff, 307ff, 323ff, 387 tectonic12, 26, 28, 34, 66, 69, 94, 102, 106, 114, 145, 148, 197, 200, 220, 278ff, 291ff, 307ff, 335, 339 terrains 293 thickness 53, 108, 109, 129, 181, 300, 309, 317, 321 tongue (s) 269ff, 293 volume 100, 259, 262, 317 wall 30, 71ff, 103, 114, 243, 308ff, 325 weld 293, 297 wings 303 Salt basins of Brazil 266 Salt basins of the world 255ff Salt basins of West Africa 261, 266 Salzwedel 32, 438 San Andreas Fault System 5 San Juan Volcanic Field 174 Sandstone 30, 112, 133ff, 159ff, 185ff, 219ff, 325, 365, 368ff, 394ff, 418 Scandinavia 19, 59, 120, 159, 164, 189, 229, 231, 233ff, 357 Schleswig-Holstein 145, 185, 237ff, 418 Sea Floor 65, 161, 245 Sea-level eustatic 3, 8, 128, 161, 167, 169, 187, 209, 210, 230 global 5, 8, 26, 130, 161ff, 186ff, 223 rise 128, 161ff, 186ff, 219ff Seal (s) 9, 10, 31, 207, 256, 325ff, 367ff, 389, 400, 410, 411, 417ff, 440, 445 Seal leakage 373ff Seawater 28, 171, 187, 251ff, 348ff, 389ff Seawater chemistry 253

Subject Index Sedimentation mode of 169 rate 28, 48, 148, 165ff, 209, 210, 293 Sediment(ary) budget 170, 218 cannibalism 170 cover 26, 29, 37, 74ff, 104ff, 213ff, 302 cycle(s) 10, 28, 157ff, 204, 435 fill 3ff, 37, 38, 50, 61 73, 89, 97, 125, 169, 213, 266, 304, 339, 340 infill 9, 19, 115 loading 24, 48, 297 sequence 8, 48, 53, 77, 78, 112, 146, 165, 204, 210, 214, 228, 297ff, 339, 367, 388, 398, 439ff supply 5, 72, 161, 169, 204, 205, 218ff system 8, 9, 134, 141, 173, 262, 335, 347, 367, 383, 426 volcanoclastic 176 Sediment facies Lower Rotliegend 159, 176ff, 181, 311 Upper Carboniferous 145, 166, 179, 433, 437 Upper Rotliegend I 159, 178, 179, 181 Segeberg-Plön Salt Wall 243 Seismic anisotropy 72 attributes 68, 413 data 3, 18, 19, 48, 68ff, 106ff, 129, 130, 173, 190, 292ff, 308ff, 411 interpretation 68, 72, 98, 108, 292ff, 324, 325 line (s) 18ff, 48, 71, 80, 107ff, 161, 226, 228, 240, 243, 268, 269, 312ff, 326 method 37, 67ff processing 68ff reflection 18, 20, 48, 53, 72ff, 130 reflectors 88, 266, 293ff refraction 79,81 sections 48, 68ff, 130, 172, 293ff, 309ff stack 69, 196, 264, 293 tomography 19, 69, 81 variance analysis 302 wide-angle refraction 6 Seismicity 7, 50ff, 119ff, 236, 241 Seismotectonics 236 Shale 31, 32, 83, 91, 133ff, 166, 169, 186ff, 325ff, 356, 365, 368ff, 390, 407, 415ff, 438ff Shale-membrane filtration 390 Shear zone 20, 45, 46, 83, 111 Shorelines 169, 233 Shortening 48, 99ff, 212ff, 268, 269, 293ff SHRIMP dating 176 Silesian-Moravian Gate 159, 191ff Sill complexes basaltic to andesitic 176 Silurian 91, 21, 92, 94, 261 Simple shear 44ff, 384 Slickensides 98, 117 Slide 188, 193, 218 Slope Failure 218, 236 Slump 188, 193, 216ff Smectite 368ff, 390, 391, 407, 427, 441 Soil 28, 79, 220, 233, 337, 338, 447 Sole Pit Basin 25, 313 Solid bitumen reflectance 138ff Solution-precipitation creep 277ff Sorgenfrei-Tornquist Zone (STZ) 19ff, 57, 80ff, 125ff, 221, 308, 309

Source area 8, 168ff, 206, 219ff, 421 point 85, 94 rock 8ff, 28ff, 137ff, 166, 168, 201, 203, 227, 324ff, 359ff, 383, 386, 400ff, 411ff, 433ff South Atlantic 201, 254 Southern Permian Basin (SPB) 17ff, 42ff, 68ff, 108, 126ff, 159, 173ff, 182ff, 291, 307ff Space problem 38, 44 Spherical 5, 55ff, 84 Spore colour 138ff Spree 233 Steady-state creep 52 Steer’s head 44 Stephanian 6, 43, 162, 176, 408 Strain elastic 97 localization 6ff partitioning 109, 122 permanent 97 softening 46 Stratigraphy 8, 27, 68, 135, 144, 163, 164, 175, 178, 186ff, 223, 250, 266, 274, 304, 305, 352, 413, 445 Strelasund Depression 178 Strength envelope 53 Stress axes 62, 111ff boundary 64, 66 coefficient 56 exponent 56, 282, 284 far field 130, 133 field 5, 7, 34, 48ff, 77, 78, 97ff, 126ff, 161, 169, 301, 306, 307ff intra plate 47ff, 130 inversion 117 palaeo 126, 132 perturbations 109, 181 ratio 119 regime 4, 22, 48, 59, 64, 121, 206, 293, 306 relaxation 212, 279 state 48, 97, 328 state, triaxial 118 tensor, reduced 117 Stretching asymmetrical 46 discontinuous 42, 43 factor 39ff instantaneous 39ff non-uniform 42ff Striae 117 Strike-slip 4ff, 24, 64, 95, 99ff, 159ff, 175, 179, 197, 206, 211, 265, 292, 293 Stromatolite 159 Structural balancing 296, 297, 324, 326 control 71, 125, 128, 292 decoupling 299, 301 environments 293 geometry (ies) 98ff, 292 regimes 291 restoration 98ff, 297, 306, 325 Shortening 48, 99ff, 268, 269, 293ff Sub-basin(s) 3, 5, 19, 25, 34, 37, 66, 89, 102ff, 125ff, 170, 171, 181ff, 211ff, 264, 304, 307ff, 433 Subduction zone 37, 51, 136

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Subject Index zones-temperature 137 Subhercynian Basin 110ff, 129ff, 170, 221ff, 313 Sub-salt basement 132, 293ff, 308, 311 extension 299 faults 293, 299 imaging 293 Subsidence centre 5, 25, 64, 78, 318 history 34, 41ff, 322, 444 initial 39ff isostatic 5, 233 post-rift 42, 44, 128 salt-induced 26 syn-rift 41 tectonic 37ff, 99, 101, 318ff thermal 5, 34, 39ff, 74, 121, 128ff, 159, 162, 181ff, 269, 297, 309, 310 Sudetes 19 Sulfate 187, 191, 301, 347ff, 389ff, 426, 427, 443ff Sulfate reduction bacterial 358, 401, 447ff thermochemical 399, 401, 427, 443ff Sulfide 360ff, 396, 401, 433ff Sulfur 357ff, 395ff, 447ff Supra-salt extension 300ff faulting 299ff Rotliegend 234 Susceptibility 79ff Sveconorwegian Orogeny 78, 89 Sweden 56, 60, 81, 121, 196, 219, 236, 434 Swell 103ff, 131, 163ff, 193ff, 218ff, 304, 308ff Sylvite 254, 277, 349ff Syn-rift 41, 42, 74, 201, 202 Syntectonic deposition 219 Syntectonic strata 106 T Tectonic event 12, 24, 29, 34, 94, 109, 113, 125, 129, 162, 167, 172, 185, 197, 210, 250, 318, 319, 446 loading 56ff overprint 173 stress 48, 55, 99, 119ff, 228, 235ff, 313, 319, 348 subsidence 37ff, 99, 101, 318ff thin-skinned 77, 94, 106, 107, 318 Tectonic Setting of halite basins 262, 266 Teisseyre-Tornquist Zone 26, 102, 126, 128, 179, 309 Tellingstedt 239 Temperature anomaly 323, 456 distribution 37, 41, 42, 133, 137, 146, 340, 341, 358 disturbance 323, 330, 337ff field 61, 125, 133ff, 323ff, 387, 388 global 157, 159 gradient 39, 41, 66, 133, 257, 324, 332, 335, 341ff history 136, 326, 430 palaeo 133ff, 159, 449 parameters 137, 141 surface 39, 133ff, 157, 159, 256, 323, 431 Tempestite 8, 192, 219 Tensional stress 48, 128, 130, 187, 200, 244 Terrain 37, 275, 293, 391 Terrigeneous 168, 171

Tertiary 25, 48, 49, 58, 65, 73, 92, 105ff, 131, 132, 143ff, 159, 166ff, 181, 210, 218, 228, 231, 233, 245, 258, 272, 273, 296, 301ff, 308ff, 325ff, 352, 357, 376ff, 400, 433, 438, 444, 446 Tethys 6, 28, 130, 159ff, 184ff, 259, 261 Thermal conductivity 30, 133, 241, 323, 327, 341ff, 432 dome (ing) 65, 66, 306 expansion 38ff, 333ff history 41, 57, 142, 221, 391, 397, 398, 438 maturation 138, 140, 324, 422, 439, 440, 449ff modelling 140, 400 overprint 173ff, 394, 443 perturbation 38, 39, 61 relaxation 24, 43, 162 subsidence 5, 34, 39ff, 74, 121, 128ff, 159, 162, 181ff, 269, 297, 309, 310 Thermochronology 142, 149, 168, 227, 391, 397, 398 Thermodynamic 348, 367, 390, 402ff, 442ff Thermohaline convection 332ff, 386, 387 flow 333 Thickness distribution 126ff, 200, 223, 224, 299, 300, 307ff, 413 Thin sheet 55ff Thin skinned 45, 77ff, 106ff, 268, 293ff, 308ff Thinning 4, 5, 20, 38ff, 82, 90, 106, 115, 149, 162, 172, 174, 205, 293, 310ff Thrust-load 212 Thrust-load basin 212 Thuringian Forest 173, 175, 219 Till 233, 245 Time-lapse 68 Tithonian 167, 199ff Toarcian 31, 32, 166, 201ff, 415, 416 Top seal 372ff Topography(ic) low 3, 5, 303 TOR 19, 20, 60, 80, 81, 428 Tornquist Fan 24 Tornquist Ocean 72, 79, 80 Tornquist-Teisseyre Zone 19, 20, 25, 103 Tornquist Zone (see also Sorgenfrei-Tornquist and Tornquist-Teisseyre Zone) 5, 19ff, 57, 60, 80ff, 102ff, 125ff, 179, 189, 221, 308, 309 Touchdown structure 71, 94 Trans-European Suture Zone 19, 79ff, 231 Transgression(ive) 28, 72, 128, 143, 159ff, 183ff, 220ff, 304, 330, 400 Transient creep 52 Transport properties 277ff, 344, 372, 381 Transpression(al) 5, 13, 72, 211, 274 Transpressive 105, 131, 184, 306, 313 Transtension(al) 5, 13, 24, 30, 37, 104, 121, 122, 126ff, 162, 173, 184, 194, 206, 210, 313 Triassic late 28, 64, 72, 77, 78, 104, 114, 122, 129, 130, 143ff, 158, 166ff, 181, 196ff, 202, 254 259, 270, 296ff, 309ff, 352, 407ff early 42, 75, 104, 129, 141,157, 159, 163ff, 171, 219, 221, 228, 254, 259, 299, 302, 307ff, 338, 352, 365, 447, 448 middle 18, 77, 91, 118, 159, 169, 171, 194, 198, 296, 302, 311, 320, 338, 358 salt 28, 251, 266ff, 303, 312 Triassic salt basin 266, 267 Tunnel Valleys 233, 240ff Turbidite 8, 12, 13, 193, 218 Turonian 66, 112, 119, 131, 158ff, 210, 216ff

Subject Index U U/Pb age 177, 182, 189, 200 Uchte area 151 Unconformity(ies) 5, 28, 47, 48, 65, 77, 78, 98, 108, 115, 130ff, 159ff, 181ff, 213ff, 244, 266, 269, 293ff, 311, 315, 319 Underplating 5, 94 Uniform stretching 38ff Unroofing sequence 222, 227, 228 Uplift 4ff, 26, 28, 38ff, 72ff, 94, 98ff, 125, 130ff, 142ff, 159ff, 184ff, 211ff, 234ff, 258, 300ff, 311ff, 331, 386, 392, 394, 398, 401, 417, 418, 432, 444, 446, 458 Uplift Rates 161, 215ff, 237, 243 Up-doming 4, 306 Upper Rotliegend II 158, 159, 173ff, 181ff Upwelling 38, 39, 45, 255ff, 271, 272, 414 Upper Rhine Valley (Oberrheingraben) 4, 136 Urstromtäler 29 V Valanginian 167, 207ff, 318 Variscan 19ff, 48, 61, 63, 75, 80ff, 111, 112, 121, 125ff, 149, 160, 162, 169, 173ff, 181, 187, 192, 235ff, 249, 261, 267, 444 Variscan Deformation Front 80ff, 128, 175, 179, 187, 235 Variscan foreland 162 Variscan Orogen (y) 20, 94, 121, 125, 149, 169, 173, 261, 444 Variscan thrust and fold belt 179 Variscides 20, 22, 89 Vein hydrothermal 111, 113, 395 Viking Graben 45, 65, 79, 159, 185ff, 407, 415, 416 Viscoelastic Response 235 Viscosity 55ff, 214, 254, 332, 333, 359, 361, 369, 384, 427 Vitrinite 138ff, 325, 361, 364, 397, 398, 414ff, 433ff Vitrinite reflectance 138ff, 325, 361, 364, 397, 398, 414ff, 433ff Volatile matter yield 139, 140 Volcanic 4, 8, 24, 103, 126, 127, 136, 142, 146, 159ff, 173ff, 183, 201, 206, 209, 254, 266, 273, 274, 356, 368, 390, 395, 403, 408, 416, 434, 443 Volcanic (rocks) basaltic 173ff facies 174ff gamma ray log 177 immobile elements 177 Volcanic succession geochronology 175 stratigraphy 175ff Volcanism topography forming 177, 180 Volcano-topographic effects 180

W Wallace-Bott hypothesis 117 Water chemistry 348ff, 389ff compaction 348 connate 348, 358, 388, 398 content 139, 140, 285 formation 134, 347ff, 388, 390ff, 452, 456, 458 fresh 355, 357, 390, 413 meteoric 9, 134, 192, 251, 264, 276, 348ff, 386ff, 389, 391, 399ff, 426, 434, 453 saline 17, 264, 336ff, 348ff, 357, 390, 398, 408 thermo-mineral 357 thermo-saline 348 Water-rock interaction 348ff, 389ff Wealden 17, 28, 146, 148, 153, 159, 167, 209, 210, 385, 415 Wealden coals 17 Weathering 168, 177, 180, 220ff, 228, 270, 275, 348ff Weichselian Glaciation 233 Weser 103, 104, 122, 158ff, 193ff, 233, 237, 300, 303, 433 Weser trough 103, 104, 163 West Brandenburg High 174, 180 West Netherlands Basin 143, 149, 153, 165, 200 Westphalian 22, 24, 28, 30, 31, 85, 163, 436ff Wetting 374, 380, 385 Wiehengebirge 151 Wietze oil field 18 Williston Basin 259ff, 419 Wilson Cycle 4, 13, 258, 265 World stress map (WSMP) 57, 61, 62, 118, 121 Wrench tectonic intracontinental 173 X Y Z Zagros foreland 261, 265 Zandvoort Ridge 143 Zechstein Basin 29, 249ff, 308, 328, 445 evaporates 173, 400 Group 163, 185ff, 249 palaeogeography 128, 186 salt 17, 31, 53, 58, 71, 90, 104ff, 128, 190ff, 224, 249ff, 291ff, 307ff, 326 stratigraphy 250 Zircon fission tracks 149

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