BASIN ANALYSIS

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Basin Analysis

Principles and Applications

Second Edition

Philip A. AllenDepartment of Earth Sciences,

ETH Zürich

John R. AllenBHP Billiton,

Australia

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© 1990, 2005 by Blackwell Science Ltda Blackwell Publishing company

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First edition published 1990Second edition published 2005 by Blackwell Publishing Ltd

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Allen, P. A.Basin analysis : principles and applications / Philip A. Allen and John R. Allen. – 2nd ed.

p. cm.Includes bibliographical references and index.ISBN 0-632-05207-4 (pbk. : alk. paper)

1. Sedimentary basins. 2. Petroleum–Geology. I. Allen, John R. (John Richard), 1953– II. Title.

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Preface to second edition vii

PART 1 THE FOUNDATIONS OF SEDIMENTARY BASINS 1

1 Basins in their plate tectonic environment 3

Summary 31.1 Compositional zonation of the Earth 41.2 Rheological zonation of the Earth 71.3 Plate motion 81.4 Classification schemes of sedimentary

basins 12

2 The physical state of the lithosphere 20

Summary 202.1 Stress and strain 212.2 Heat flow: conduction and convection 312.3 Gravity and isostasy 452.4 Rock rheology 51

PART 2 THE MECHANICS OF SEDIMENTARY BASIN FORMATION 61

3 Basins due to lithospheric stretching 63

Summary 633.1 Introduction to rifts, failed rifts, and

passive continental margins 64

Contents

3.2 Geological and geophysical observations in regions of continental extension 65

3.3 Introduction to models of continental extension 73

3.4 Uniform stretching of the continental lithosphere 77

3.5 Modifications to the uniform stretching model 85

3.6 A dynamic approach to lithospheric extension 97

3.7 Mantle plumes and igneous activity associated with continental extension 106

3.8 Estimation of the stretch factor and strain rate history 112

4 Basins due to flexure 116

Summary 1164.1 Basic observations in regions of

lithospheric flexure 1194.2 Flexure of the lithosphere: geometry

of the deflection 1244.3 Flexural rigidity of the oceanic and

continental lithosphere 1304.4 Lithospheric buckling 1394.5 The dynamics of orogenic wedges 1444.6 The modeling of foreland basins 155

5 Effects of mantle dynamics 167

Summary 1675.1 Fundamentals and observations 1685.2 Dynamic topography 175

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vi CONTENTS

6 Basins associated with strike–slip deformation 190

Summary 1906.1 Overview 1916.2 The structural pattern of strike–slip

fault systems 2016.3 Basins in strike–slip zones 207

PART 3 THE SEDIMENTARY BASIN-FILL 219

7 The sediment routing system 221

Summary 2217.1 Introduction 2227.2 Weathering 2227.3 Terrestrial sediment and solute

yields 2277.4 Measurements of erosion rates 2417.5 The functioning of sediment routing

systems 245

8 Basin stratigraphy 266

Summary 2668.1 A primer on process stratigraphy 2688.2 Stratigraphic cycles: definition and

recognition 2758.3 Driving mechanisms for stratigraphic

patterns 2868.4 Numerical simulation of stratigraphy 3068.5 Depositional systems 3118.6 Relation of depositional style to basin

setting 327

9 Subsidence and thermal history 349

Summary 3499.1 Introduction to geohistory analysis 3519.2 Porosity loss during basin subsidence 3519.3 Subsidence history and backstripping 3599.4 Introduction to thermal history 3669.5 Theory: the Arrhenius equation and

maturation indices 3679.6 Factors influencing temperatures and

palaeotemperatures in sedimentary basins 368

9.7 Measurements of thermal maturity in sedimentary basins 378

9.8 Application of thermal maturity measurements 391

9.9 Geothermal and paleogeothermal signatures of basin types 395

PART 4 APPLICATION TO PETROLEUM PLAY ASSESSMENT 403

10 The petroleum play 405

Summary 40510.1 From basin analysis to play concept 40610.2 The petroleum system and play

concept 40710.3 The petroleum charge system 41410.4 The reservoir 45210.5 The regional topseal 46710.6 The trap 474

References 495Index 532

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The first edition of Basin Analysis was written in 1988 and1989 and published in 1990. A decade and a half havetherefore passed since the writing of the first edition.Fifteen years is a long time between any two editions ofa popular textbook, but in the field of basin analysis it isan eternity. The field has exploded with new data, newtechnology, and new concepts. This second editionattempts to take account of these new developmentswhile retaining the successful basic structure and philo-sophy of the original text.

We originally defined basin analysis as the integratedstudy of sedimentary basins as geodynamic entities.This is even more the case today. It is now realized thatthe geodynamic system involving basin developmentmust also include the processes of rock exhumation,erosion and sediment transport in the source terrains ofsedimentary basins. In other words, the sediment supplyto a basin is more than simply a boundary condition. Thesediment efflux of eroding catchments is a thread linkingintimately coupled process systems governed by a wholespectrum of driving mechanisms. Tectonics affect catch-ment topography as well as basin subsidence. Climateaffects weathering and erosion as well as depositionalenvironments and sedimentary products. Base-levelchange affects river incision and sediment transport aswell as accommodation in the basin. To understand abasin as a geodynamic entity therefore requires an appreciation of the coupling between mantle, litho-sphere, oceans, atmosphere, and land surface. The per-spective is expansive and multidisciplinary. Basin analysisis no more the preserve of the geophysicist than the strati-grapher. It is the classic integrative topic in the curric-ulum and in our research portfolios.

The second edition of Basin Analysis is a major updat-ing and reassessment of the first edition. We have

retained the basic structure of first considering the fun-damental physical state of the lithosphere before dis-cussing the two grand classes of lithospheric deformationrelevant to basin development – stretching and flexure.We have added a chapter on the effect of mantle dynam-ics, since it is now generally accepted that flow in themantle related to the subduction of oceanic slabs, plumeactivity, and the slow incubation of heat beneath super-continental assemblies impinge on basin development.Basins related to strike–slip motion continue to be treatedin their own chapter because of their specific, intimatelinkage to deformation within the principal displacementzone. The chapters on the basin-fill have been radicallyrevised. A new chapter on the erosional engine for sedi-ment delivery to basins has been included to reflect themassive and exciting advances in this area in the lastdecade. The discussion of basin stratigraphy highlightsits structure or architecture and proposes explanations interms of the external driving mechanisms or of theunforced, internal dynamics responsible. The techniquesnow used routinely in approaching problems in basinanalysis have much expanded. We have updated subsi-dence analysis and measurements of thermal maturity oforganic and nonorganic components of the basin-fill asthe staple diet of basin analysts, but have included newthermochronological and exposure dating tools that haverevolutionized our understanding of the rates of denuda-tion of source regions, storage, and transfer to basins.Finally, the section on the application to the petroleumplay has been retained, condensed, and updated toinclude the important petroleum system concept. Itremains a useful summary of the scientific basis for petro-leum geology and how many of the concepts and tech-niques discussed earlier in the book can be applied in acommercial arena.

Preface to second edition

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viii PREFACE

The educational or intellectual philosophy adopted inBasin Analysis is that it is necessary to build up under-standing by examining the underlying principles for theway in which basins function, rather than by presentinga blizzard of detail on the variability of the real world.We do not mean to be sectarian in the deductive versus

inductive debate. The messy ground-truth of sedimen-tary basins, their structure, thermal and subsidencehistory, and stratigraphic architecture, is that eachexample has its own set of attributes, or to be anthropo-morphic, its “individuality.” But it is not an easy task toinduct the functioning of basins from such a blizzard ofdetail. We therefore present wherever possible a simpli-fied but useful physical analysis, backed up where appro-priate with elementary calculus, as the means to enhanceour understanding of basins generically. A further advan-tage of this approach is that students become armed witha wider range of scientific tools that can be used in aspectrum of other Earth Science and non-Earth Scienceapplications. This is important in our current climate of fostering “transferable skills.” The use of this first-principles approach is necessarily uneven in the text.Where a mathematical derivation is an optional extra, itis boxed separately from the rest of the text so as not tointerrupt the flow. In most cases, however, the student willbenefit from working through the simple mathematicalanalysis, inserting typical parameter values, and derivinga quantitative result. We have tried to encourage this“interactiveness” by providing practical examples withsolutions on freely accessible web pages at ETH Zürich (www.erdw.ethz.ch/Allen).

Our perspective on developments over the last 15 yearsis that some hot issues have become luke warm, whereasothers that were previously invisible or dormant havebecome widely debated. Such is scientific progress. As anexample of the former, the art of sequence stratigraphy,with its controversial overemphasis on eustatic sea-levelchanges as the driving forces for stratigraphic packaging,has matured into a valuable observational tool with inter-pretation founded on a far better appreciation of thedynamics of accommodation generation and sedimentsupply. As an example of the latter, numerical models ofa multilayer lithosphere now provide new insights into arange of problems involving the dynamics of lithosphericcontraction and extension. A forceful and exciting com-munity of scientists is simultaneously providing keyobservations and models to enable such improved solid-Earth models to be coupled to basin development via thecritical interface of the Earth’s surface. Such a successfulmerging of geomorphology, geophysics, isotope geo-

chemistry, sedimentology, and structural geology sends a powerful message of the value of multidisciplinaryapproaches in the Earth Sciences.

The book is intended for undergraduates who havedone the prerequisite set of introductory courses in theEarth Sciences. Those who have emerged from a soundphysical sciences background will find this book more totheir liking, but those entering Earth Science from the lifeor environmental sciences should not despair. The bookshould also prove to be useful to postgraduate studentseither as a text for taught courses in basin analysis andpetroleum geology, or as a synthesis for doctoral studentsgrappling with research problems across a wide range ofEarth Science subdisciplines.

The writing of a textbook is a long, difficult road totread. Goodness knows why anyone volunteers to do it.Any attempt at “completeness” has to be abandoned atan early stage. First, completeness can mean vast docu-mentation of trivia. Second, attempts at completenessgenerally end in noncompletion. The book project thenbecomes a never-ending introspective diary. So to theacademic staff using this book, please be understandingif sufficient credit is not given to your work. To thestudent, please do not think that you should not be visit-ing the library for supplementary reading.

Apart from primary sources in research papers, thereare a number of textbooks and compilations or reviewsthat we have found valuable to consult and use. In thegeneral area of basin analysis is Ingersoll and Busby’svery useful compilation Tectonics of Sedimentary Basins

(1995). Einsele’s Sedimentary Basins (second edition, 2000),despite its title, is strongly focused on sedimentarygeology. Leeder’s Sedimentology and Sedimentary Basins

(1999) is a stimulating text whose subtitle “From Turbu-lence to Tectonics” perfectly reflects its contents. Theedited textbook by Reading (Sedimentary Environments, thirdedition, 1996) continues to be the best comprehensivesource of all matters sedimentological. In the area ofsurface processes, useful material can be found in Sum-merfield’s Global Geomorphology (1990) and its reincarna-tion as a multiauthor compilation in Geomorphology and

Global Tectonics (2000). Allen’s Earth Surface Processes (1997)provides a physically based approach to sediment routingand depositional processes, and Burbank and Anderson’sTectonic Geomorphology (2001) trail-blazes the importantarea of neotectonics and geomorphology. Turning togeophysics and dynamics, there is little competition forthe mighty Geodynamics by Turcotte and Schubert (secondedition, 2002). Very valuable material is also found inFowler’s The Solid Earth (1990) and Isostasy and Flexure of

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the Lithosphere by Watts (2001) is very helpfully focused onsome of the key issues underpinning basin development.Diagenesis: A Quantitative Perspective by Giles (1997) is a tour

de force in the area of compaction and diagenesis of thebasin-fill. Magoon and Dow’s edited volume The Petroleum

System (1994) is a valuable addition to the literature in thefield of petroleum geology.

The journal Basin Research has published a number of important thematic sets that provide ideal huntinggrounds for students of basin analysis: some examples are Foreland Basins edited by Allen, England, Watts andGrotzinger (1992), Cratonic Basins (1994) edited by Allen,Watts, Grotzinger and Royden, Supradetachment Basins

edited by Friedmann and Burbank (1995), Sediment Supply

to Basins edited by Hovius and Leeder (1998), Interactions

of Tectonics and Surface Processes in Landscape Evolution editedby Burbank and Pinter (1999), Processes and Controls in the

Stratigraphic Development of Extensional Basins edited byGupta and Cowie (2000), and Numerical and Physical Ex-

perimental Modelling of Stratigraphy edited by Burgess andPaola (2002). Readers are also referred to the compila-tion of papers entitled Tectonics and Topography, edited byEllis and Merritts, in the Journal of Geophysical Research

(1994), and Quantitative Dynamic Stratigraphy (1990) editedby Cross, which together kick-started many quantitativeapproaches to the coupling of tectonics, surfaceprocesses, and stratigraphy.

We are grateful for the help of many colleagues, espe-cially those who found time to look at drafts of bookchapters, including Alex Densmore and Nick Richardson(ETH Zürich), and Peter Burgess (Shell, Rijswijk). AlexDensmore has cotaught a basin analysis class with PAAfor several years and has provided much helpful input.Sonia Scarselli and Chris Krugh helped to assemble thepractical exercises and solutions available on the web.Maja Reichenbach kindly provided much secretarialhelp. We are once again grateful for the support of staffat Blackwell Publishing, particularly Delia Sandford andRosie Hayden in production and editor Debbie Seymour.Finally, we dedicate this volume to our families, Carolyn,Lorna, Frances, and Richard spread across England andSwitzerland, and Annette, Michael, and Robert in Mel-bourne, Australia.

PAA, JRA, Zürich

PREFACE ix

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PA R T

1 The foundations ofsedimentary basins

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SUMMARY

Sedimentary basins are regions of prolonged subsidenceof the Earth’s surface. The driving mechanisms of sub-sidence are principally related to processes within therelatively rigid, cooled thermal boundary layer of theEarth known as the lithosphere. The lithosphere is com-posed of a number of plates which are in motion withrespect to each other. Sedimentary basins therefore existin a background environment of plate motion.

The Earth’s interior is composed of a number of com-positional and rheological zones. The main composi-tional zones are between crust, mantle and core, the crustcontaining relatively low density rocks overlain by a sedi-mentary cover. The mechanical and rheological divisionsdo not necessarily match the compositional zones.A fundamental rheological boundary is between the lithosphere and the underlying asthenosphere. The lithosphere is sufficiently rigid to comprise a number ofrelatively coherent plates. Its base is marked by a char-acteristic isotherm (c. 1330 °C) and is commonly termedthe thermal lithosphere, or alternatively mechanical lithosphere.The upper portion of the thermal lithosphere is able tostore elastic stresses over long time scales and is referredto as the elastic lithosphere. The continental lithosphere hasa strength profile with depth which suggests that a weak,ductile zone exists in the lower crust, separating a brittleupper crust and upper mantle, giving a jam-sandwichtype structure. The oceanic lithosphere, however, lacksthis low strength layer, its strength increasing with depthto the brittle–ductile transition in the upper mantle.

The relative motion of plates produces deformation,vulcanicity, and seismicity concentrated along theirboundaries, which are classified as divergent boundaries,such as the mid-ocean ridge spreading centres of the

ocean basins, convergent boundaries associated with largeamounts of shortening, such as continental collisionzones, and conservative boundaries characterized bystrike–slip deformation. Although the theory of plate tec-tonics has the premise that deformation is concentratedalong plate boundaries, the continental lithospheredeforms far from plate boundaries, and appears tobehave at geological time scales more like viscous sheetsthan as rigid stress guides.

Sedimentary basins have been classified principally interms of the type of lithospheric substratum (i.e., conti-nental, oceanic, transitional), their position with respectto the plate boundary (intracratonic, plate margin), andtype of plate motion nearest to the basin (divergent, con-vergent, transform). The formative mechanisms of sedi-mentary basins fall into a small number of categories,although all mechanisms may operate during the evolu-tion of a basin:• Isostatic consequences of changes in crustal/lithospheric thick-

ness, such as caused mechanically by lithosphericstretching, or purely thermally, as in the cooling andsubsidence of the oceanic lithosphere as it moves awayfrom oceanic spreading centres;

• loading (and unloading) of the lithosphere causes a deflec-tion or flexural deformation and therefore subsidence(and uplift), as in foreland basins;

• viscous flow of the mantle causes nonpermanent subsidence/uplift known as dynamic topography.From the point of view of lithospheric processes there

are two major groups of basins: (i) Basins due to litho-spheric stretching, belonging to the rift-drift suite, and (ii)basins formed primarily by flexure of continental andoceanic lithosphere.

C H A P T E R

1 Basins in their plate tectonic environment

Assumptions, hasty, crude and vain,

Full oft to use will Science deign;

The corks the novice plies today

The swimmer soon shall cast away

(A.H. CLOUGH, POEM (1840))

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4 CHAPTER 1

Inspection of any map showing hydrocarbon occur-rences (e.g., St John et al. 1984) reveals their clusteredpattern. In general, provinces of hydrocarbon occur-rence correspond to the locations of sedimentary accu-mulations greater than about 1 km thick. Theseaccumulations include sedimentary basins in the strictsense, implying zones of pronounced subsidence (Ballyand Snelson 1980) but also carbonate bank build-ups on elevated oceanic crust, cratonic arches, and so on whichbecome fossilized in the geological record. However,sedimentary basins located at sites of prolonged and substantial subsidence are of overriding importance.

Historically, basin studies have developed from a numberof distinct viewpoints such as that of stratigraphicsequences and their relation to sea-level fluctuations (Sloss1950, 1963), the geosyncline (Kay 1947, 1951; Aubouin1965), and, more recently, the concept of plate tectonics(Dickinson 1974; Ingersoll 1988). The location of sedimen-tary basins and their driving mechanisms are intimatelyassociated with the motion of discrete, relatively rigid slabs,which together represent the cooled thermal boundarylayer of the Earth, and with the convective flow of theunderlying mantle. The outer shell of the Earth comprisesa relatively small number of these thin, relatively rigidplates, and they are in a state of motion with respect to eachother. Such motions set up plate boundary forces that maybe transferred considerable distances into the interior of theplates, so that sedimentary basins exist in a backgroundenvironment of stress set up by plate motion. The lithos-pheric plates are the surface manifestation of a slow thermalconvection in the mantle, and are subject to differentialthermal stresses along their bases. The mantle and litho-sphere therefore do not operate as independent systems. Wesee spectacular evidence for the interaction of mantleprocesses and the lithosphere in the volcanic and topo-graphic expression above plumes that have risen from thecore–mantle boundary. We also discern, though less spec-tacularly, the effects on mantle flow of the subduction ofcold slabs of oceanic crust at ocean–continent boundaries.

Some basic ideas on plate tectonics and Earth struc-ture are introduced in this chapter in so far as they helpto explain the location and evolution of sedimentarybasins. More exhaustive summaries can be found inWyllie (1971), Cox (1973), Le Pichon et al. (1973), Smith(1976), Bird (1980), Cox and Hart (1986), and Keareyand Vine (1996). Ingersoll and Busby (1995) is a collec-tion of individually written chapters of relevance to basinanalysis and plate tectonics.

The Earth’s interior is composed of a number ofessentially concentric zones that are defined on the basis

of either compositional changes, or mechanical/rheological changes.

1.1 COMPOSITIONAL ZONATION OF THEEARTH

There are three main compositional units; the crust,mantle, and core (Fig. 1.1).

1.1.1 Oceanic crust

The crust is an outer shell of relatively low density rocks.The oceanic crust is thin, ranging from approximately 4to 20 km in thickness, 10 km being “normal,” and withan average density of about 2900 kg m-3. It comprises anumber of layers that reflect its mode of creation: anupper veneer (layer 1) of unconsolidated or poorly consolidated sediments, generally up to 0.5 km thick; anintermediate layer 2 of basaltic composition, consistingof pillow lavas and associated products of submarineeruptions; and a layer 3 of gabbros and peridotites thatmay form the parent rocks which upon differentiationgive rise to the basalts of layer 2. The oceanic crust hasbeen thought to be distinctly layered in terms of velocityof seismic waves, but more recent views are that it pos-sesses a more gradual and continuous increase in veloc-ity with depth.

The lifetime of oceanic crust is short, despite the factthat it occupies about 60% of the surface of the Earth(c. 3.2 ¥ 109 km2). This is because as the oceanic crustcools during aging it becomes gravitationally unstablewith respect to the mantle; as a result it is consumed. Thisexplains why the oldest oceanic crust in today’s oceans isas young as Jurassic in age (c. 150 Ma). Compared to thecontinents, the oceanic crust therefore has a very shortlifespan.

1.1.2 Continental crust

The continental crust is thicker, ranging from 30 to 70 km, but with an “average” thickness of perhaps 35 km.It was originally thought to be divided into two layers,each with a distinct composition and density: (i) An upperlayer with physical properties similar to those of granites,granodiorites, or diorites overlain by a thin veneer ofsedimentary rocks. This so-called “granitic layer” has a thickness of between 20 and 25 km and a density of 2500–

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2700 kg m-3. The term “granitic” is, however, misleading,since average densities are greater than that of granite;(ii) a lower layer of primarily basaltic composition, butthe pressure and temperature at depths in excess of

25 km imply that the rocks are granulites, or their highpressure, high temperature equivalents, eclogites oramphibolites. The density of this lower layer is2800–3100 kg m-3. These layers may not in reality be welldefined, and instead a more continuous variation of com-position with depth may exist.

Information on the density of crustal rocks has beenobtained largely by observations on seismograms of thespeed of seismic waves passing through the variouslayers, coupled with laboratory experiments on rockmaterials. The existence of a low velocity crust was dis-covered by the geophysicist Mohorovicic shortly after theturn of the century. At the crust–mantle boundary,seismic P (longitudinal) wave velocities increasemarkedly; this abrupt increase in velocity may reflect acorresponding increase in rock density (Fig. 1.1). Thishorizon is known as the Mohorovicic discontinuity orMoho. The Moho varies in depth considerably. The con-tinental crust thickens under orogenic belts, thins underzones of rifting, and attenuates completely at continen-tal margins (Fig. 1.2).

In some regions, particularly the attenuated margins of continents, the crust is intermediate in character andthickness between typical oceanic and continental varieties. This may be due to the injection of dense intrusions, to metamorphism, or to other processesaccompanying stretching. In particular, the depth to theMoho may be abnormally great due to the igneousunderplating of the crust during plume activity. Severalkilometers of underplate emplaced in the Early Tertiaryover the head of the Iceland plume have been interpretedfrom seismic experiments carried out on the northwestEuropean continental margin.

1.1.3 Mantle

The mantle is divided into two layers: the upper and lowermantle. The upper mantle extends to about 680 km ±20 km and is punctuated by phase transitions. The innermantle extends to the outer limit of the core at 2900 km,with an increasing density with depth.

Although there are of course no in situ measurementsof the composition of the mantle, it can be estimatedfrom the chemistry of volcanic and intrusive rocksderived by melting of the mantle, from tectonicallyemplaced slivers of mantle rock preserved in orogenicbelts known as ophiolites, from nodules preserved in vol-canic rocks, from minerals brought to the surface explo-sively in kimberlites, and importantly, from the remote

BASINS IN THEIR PLATE TECTONIC ENVIRONMENT 5

ContinentOcean

Moho

Low velocity zone

Phase boundary

Asthenosphere

Mesosphere

0 km

500 km

400 km

680 km

2900 km

5100 km

6378 km

0 km

Lower transition zone

Lithosphere

Core"D"

Crust Mantle lithosphere

(a)

(b)

Vp

Upper transition zone

Upp

erm

antle

Low

er m

antle

Out

erco

re

Inne

rco

re

[olivine]

[spinel]

[perovskiteand MgO]

8.0 8.5 9.0 km s–1

Fig. 1.1 The main compositional (a) and rheological (b)boundaries of the Earth. The most important compositionalboundary is between crust, mantle, and core. There are strongcompositional variations within the continental crust and com-positional variations caused by phase changes in the mantle. Vp

is velocity of P wave. The main rheological boundary is betweenthe lithosphere and the asthenosphere. P wave velocities increasemarkedly beneath the Moho, but decrease in a low velocity zonerepresenting the weak asthenosphere. The lithosphere is rigidenough to act as a coherent plate. P wave velocity in (a) fromwestern Europe after Hirn (1976).

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(a)

(b)

Contours in km below sea level

DEPTH TO MOHO

Fig. 1.2 Depth to the Moho below sea-level from (a) Europe and (b) North America, after Allenby and Schnetzler (1983) and Meissner (1986). In Europe, thick continental crust occurs in the Pyrenean–Alpine–Carpathian orogenic belt and in the Scandinavian Shield. Thin continental crust occurs along the Atlantic margin and in regions of continental stretching, such as theNorth Sea, western Mediterranean, Pannonian Basin, and Black Sea. In North America, thick continental crust is associated withthe batholiths of the Sierra Nevada region, the Appalachian orogenic belt, the American midcontinent and the Canadian–NE USAshield. Thin crust is associated with the plateau basalts of the Snake River area, sites of rifting such as the Mississippi Embayment,and along the western active continental margin, eastern passive continental margin and the Gulf of Mexico. NS, North Sea; SC,Scandinavian Shield; PY, Pyrenees; AP, Apennines; PB, Pannonian Basin; H, Hellenides; SR, Snake River; SN, Sierra Nevada; ME,Mississippi Embayment; AP, Appalachians; GM, Gulf of Mexico. Reproduced courtesy of Elsevier.

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but sophisticated measurement of the mantle usingseismic waves. The main constituent of the mantle isthought to be olivine, mostly the Mg-rich varietyforsterite.

Olivine is known to undergo phase changes to denserstructures at pressures equivalent to depths of 390–450 km and c. 700 km in the Earth. At 390–450 kmolivine is thought to change to spinel via an exothermicreaction involving a 10% increase in density. At c. 700 kmspinel changes to perovskite and magnesium oxide in an endothermic reaction. These phase changes can be recognized by changes in the velocity of S waves(McKenzie 1983) and may determine the scale of con-vection in the mantle (Silver et al. 1988).

1.2 RHEOLOGICAL ZONATION OF THE EARTH

The mechanical or rheological divisions of the interiorof the Earth do not necessarily match the compositionalzones. One of the rheological zonations of primary interest to students of basin analysis is the differentiationbetween the lithosphere and the asthenosphere. This is because the vertical motions (subsidence, uplift) in sedimentary basins are principally a response to thedeformation of this uppermost rheological zone of theEarth.

1.2.1 Lithosphere

The lithosphere is the rigid outer shell of the Earth, com-prising the crust and the upper part of the mantle. It isof particular importance to note the difference betweenthe thermal and elastic thicknesses of the lithosphere. It is generally believed (e.g., Parsons and Sclater 1977;Pollack and Chapman 1977) that the base of the lithosphere is represented by a characteristic isotherm(1100–1330°C) at which mantle rocks approach theirsolidus temperature. This defines the thermal lithosphere.Typical thicknesses of lithosphere under the oceansvaries from c. 5 km at mid-ocean ridges to c. 100 km inthe coolest parts of the oceans. The lower boundary of the lithosphere is poorly defined under continents,depths of 100 to 250 km being typical. The stepwiseincreases in velocities of S and P waves with depththrough the lithosphere suggest that it contains composi-tional boundaries within it.

The rigidity of the lithosphere allows it to behave as acoherent plate but only the upper part of the lithosphereis sufficiently rigid to retain elastic stresses over geologi-cal time scales (say 107 years). Below this upper elastic

lithosphere creep processes efficiently relax elastic stresses,so that there is a physical and conceptual differencebetween the elastic lithosphere and the thermal litho-sphere. The lithosphere below the upper elastic portionmust therefore be sufficiently soft to relax elastic stressesbut sufficiently rigid to remain a coherent part of thesurface plate.

The oceanic and continental lithosphere differ in theirstrength (Fig. 1.3). The strongest part of the oceanic lithosphere occurs in the mantle between 20 and 60 kmdepth, below which it becomes increasingly ductile. Thecontinental lithosphere, however, appears to be markedlyzoned rheologically. In particular, the upper seismicallyactive brittle zone overlies a generally aseismic zone thatmay deform by ductile processes. This mid-lower crustalductile zone has been invoked as a level of detachmentof major upper crustal faults (e.g., Kusznir and Park1987) (§3.5.2 and §3.5.3). There is a second, deeperstrong layer in the mantle part of the continental litho-sphere where earthquakes occasionally occur (Chen andMolnar 1983).

There are also heterogeneities in the mantle part ofthe lithosphere, although they are small compared withthe crust. Seismological studies of western Europe (Hirn1976) suggest a highly stratified lithosphere beneath theMoho (Fig. 1.1). In particular, a “channel” of reduced Pwave velocities has been interpreted between 10 and 20 km below the Moho. This 10 km thick layer cannot beexplained in terms of partial melting since the solidustemperature is far in excess of the actual temperature –the hydration (serpentinization) of peridotites has beenpostulated as a possible mechanism. Whatever the cause,this upper low velocity channel may serve as a zone ofdecoupling of the upper lithosphere from the lowerportion of the lithosphere when acted upon by tangen-tial tectonic forces. There are few examples, however,where a process of decoupling can be unambiguouslydemonstrated at these levels.

1.2.2 Sublithospheric mantle

The underlying region, the asthenosphere is weaker than thelithosphere and is able to undergo deformation relativelyeasily by flow. The upper part of the asthenosphere isknown as the low velocity zone where P and S wave

BASINS IN THEIR PLATE TECTONIC ENVIRONMENT 7

ABA1 11/22/04 4:31 PM Page 7

8 CHAPTER 1

transmission speeds drop markedly, presumably due topartial melting (Fig. 1.4).

Studies of minute variations in the transmission speedsof seismic waves has allowed the structure of the deepmantle to be visualized and mapped (Dziewonski andWoodhouse 1987), a topic known as seismic tomography.Zones of faster than average seismic velocity are attrib-uted to propagation through denser rock which in turnis most likely due to a cooler temperature. Zones withslower than average seismic velocity are likewise thoughtto be due to warmer temperatures. Variations in tem-perature are probably caused by large-scale convection.Far from being inert, the mantle represents a vast volumeof rock which dynamically interacts with the lithosphere.Instabilities rise from the core–mantle boundary as plumes

of hot material which impinge on the base of the over-lying lithosphere and may have a major role in contin-

ental break-up (Burke and Dewey 1973; White andMcKenzie 1988). The thermal effects of the subductionof cold oceanic lithosphere and the insulating effects ofsupercontinental assemblies are also thought to be re-cognizable in the thermal structure of the mantle (Gurniset al. 1996).

1.3 PLATE MOTION

Plate tectonics can operate because the lithosphere iscomposed of a number of coherent rheological “plates”(Fig. 1.5). The underlying concepts of relative platemotion come from studies of focal mechanism solutionsof large earthquakes and observations of the distributionof earthquake epicenters, and from studies of magneticlineations in the ocean basins. The nature and rates of

Moho

Dep

th (

km)

Frictional sliding

Brittle-ductile transition

Plastic flow

(a) Oceanic lithosphere

Maximum rock strength

Plastic flow

Plastic flow

Frictional sliding

Brittle-ductile transition

Moho

Maximum rock strength

(b) Continental lithosphere

Olivine

Quartz

Brittlefailure50

100D

epth

(km

)

50

100

Fig. 1.3 Strength profiles for the oceanic (a) and continental (b) lithosphere, based on Molnar (1988) and Sammonds (1999). Theyield strength of the continental and oceanic lithosphere is plotted as a function of depth. The olivine rheology of the oceanic lithosphere provides a strong elastic core extending to depths of over 50 km. The quartz or quartz-felspar rheology of the contin-ental lithosphere causes a weak, ductile layer at equivalent depths. A second brittle–ductile transition occurs in the mantle litho-sphere because of the compositional change to an olivine rheology. The elastic lithosphere is the upper portion that is able to storeelastic stresses over long time periods. The base of the thermal lithosphere is a mechanical boundary separating the relatively strongouter shell of the lithosphere from the very weak asthenosphere.

ABA1 11/22/04 4:31 PM Page 8

relative plate motion (Minster and Jordan 1978) governmany aspects of the geodynamic environment of basins.Increasingly, reliable information on motion within platesis provided by measurement of geodetic strains involvingthe long-term occupation of triangulation networks andthe use of ground position satellite (GPS) technology(Clarke et al. 1998).

The global pattern of seismic activity is of continuousand narrow belts of high frequency of earthquakes,bounding extensive regions of relative stability(Barazangi and Dorman 1969). The narrow zones ofearthquake activity define plate margins. Oceanic plateboundaries are very sharply defined whilst continentalboundaries are rather more diffuse.

The fact that earthquake epicenters occur at depths asgreat at 650 to 700 km along some plate boundaries sug-gests that a process exists that is capable of transferringbrittle material to depths normally associated with

deformation by flow. This process of plate subduction is responsible for both the relative youth of the oceaniccrust and the distribution of earthquake epicenters. Thefact that the interiors of plates experience only infrequentearthquake activity reflects the concentration of large relative motions of plates along their boundaries. How-ever, this does not mean that the interiors of plates donot experience significant deformation. Continentalplates clearly undergo extensive deformation a long wayfrom plate boundaries (England 1987). In such cases, thedriving force is believed to be the excess potential energyof elevated continental crust.

The lithospheric plates can be easily deformed bybending about a horizontal axis, but are highly resistantto torsion about steeply inclined axes. This latter prop-erty of strength allows the motion of plates over theEarth’s surface to be modeled assuming no internaldeformation, except at plate boundaries. But how do theoceanic and continental lithosphere compare in terms offlexural strength? Different views exist on this problem.On the one hand, oceanic plates are stronger becausethey consist of more mafic mineral assemblages, whereasthe continents contain quartz which shows ductile flowat lower temperatures than olivine (see Fig. 1.3). Theyalso contain fewer intrinsic weaknesses such as old fun-damental fault systems. On the other hand, the oceanicplates are thinner and hotter, and therefore bend moreeasily under an applied force system. Whether continen-tal or oceanic lithosphere is stronger in terms of theirresistance to bending, or flexure, is therefore controver-sial and the strength must at least in part depend onparameters such as geothermal gradient and strain rate.Continental crust, however, appears to be weaker thanoceanic crust when subjected to extensional stresses(Steckler and ten Brink 1986).

Three classes of plate boundary exist: divergent, con-vergent, and conservative (Fig. 1.6):

Divergent boundaries are typified by the mid-ocean ridgespreading centres of the ocean basins. Here, the recog-nition of magnetic bands correlated with a magneticreversal chronology (Vine and Matthews 1963; Cox1973) allows the rate of divergent plate motion to be esti-mated. Transform faults with strike–slip displacementoffset the divergent boundaries, producing a highly seg-mented pattern.

Convergent boundaries are of two classes:• Subduction boundaries where oceanic lithosphere

constitutes the downgoing plate. Ocean–ocean

BASINS IN THEIR PLATE TECTONIC ENVIRONMENT 9

Temperature °C0 500

200

400

600

800

1000

1000 1500 2000 2500

Low velocity zone Partialmelting

Adiabat

Tm Solidus melting temperature formantle material

Geotherm

Dep

th (

km)

Fig. 1.4 Variation of temperature with depth, or geotherm,and the solidus temperature for mantle material (peridotite).Where the solidus curve (Tm) and the geotherm become tan-gential, partial melting in the mantle is likely to take place,resulting in a zone of low seismic wave velocities (low velocityzone).

ABA1 11/22/04 4:31 PM Page 9

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ABA1 11/22/04 4:31 PM Page 10

boundaries, as for example, in the Mariana Islands, arecharacterized by a well-developed ocean trench andvolcanic island arcs, whereas ocean–continent bound-aries such as along the west of the Andes consist of anocean trench with an associated continental magmaticarc with intense plutonic activity.

• Collisional boundaries where continental lithosphereconstitutes the downgoing plate. Where both plates arecontinental, as in the Alps or Himalayan zones, thebuoyancy of the downgoing plate resists subduction,leading to intense and widespread deformation. Lesscommonly, oceanic lithosphere may override conti-nental lithosphere attached to subducting oceanic lithosphere, as in Taiwan.

Conservative boundaries occur where the adjoining plates aremoving parallel to each other and are therefore domi-nated by strike–slip or transform faults.

The relative movement between plates causes earthquakes, a fact demonstrated by the concentration of seismic activity along plate boundaries. Earthquakesoccur along trenches, ridges, and transforms, but they are distinctly different along the three types ofboundary:• Ridges are characterized by small to moderate earth-

quakes generated at shallow depths of <10 km;• transforms experience larger earthquakes originating

from depths of <20 km;

• subduction zones are sites of very large and deep earth-quakes, with foci occurring as deep as 700 km.

The disappearance of earthquakes at relatively shallowlevels along transforms and ridges is thought to be dueto the change in rheology from brittle (capable of storingelastic stresses before rupture) to ductile (flowing bycreep). This transition takes place in the range 600–900 °C, which corresponds to a depth of 20–30 km attransforms, but at shallower levels of about 10 km atridges where temperatures are elevated. In contrast,at subduction zones, if the plate is descending quickly itremains cool relative to its surroundings and is capableof brittle deformation to large depths. Hence, earthquakefoci along subduction zones (Fig. 1.7) may be very deep.The rate of slab subduction may affect the depth ofearthquakes, however. If the slab is only slowly sub-ducted, it may heat up sufficiently to prevent the occur-rence of earthquakes at great depths.

It is possible to study the type of motion that occurredduring a particular earthquake and to work out thestresses released at the time of the earthquake, andthereby the direction of plate motion that gave rise to thestresses. These methods are called first motion or focal

mechanism studies, since it is the first motion of the groundsurface, whether it is up/away from the source of theearthquake or down/towards the focus. Distinct radia-tion patterns result from first motion on strike–slip,normal, and reverse faults.

BASINS IN THEIR PLATE TECTONIC ENVIRONMENT 11

Mesosphere

Asthenosphere

Lithosphere Lithosphere

Transform

Ridge Subductionzone

Divergentplate boundary

Conservativeplate boundary

Convergentplate boundary

Fig. 1.6 The three types of plate boundary: convergent, divergent, and conservative (after Kearey and Vine 1996). Reproducedcourtesy of Blackwell Publishing Ltd.

ABA1 11/22/04 4:31 PM Page 11

12 CHAPTER 1

Sykes (1967) studied the first motion of earthquakesalong mid-Atlantic ridge transforms and suggested thatthe motion was strike–slip, to the east on the northernblocks and to the west on the southern blocks. This isright lateral motion – opposite to that indicated by theoffsets of the ridge – but the correct relative movementto support the plate tectonic interpretations of trans-forms as actively shearing only between the ridge seg-ments and not beyond. The earthquakes originating fromspreading centres are quite different from those beingproduced at transforms. First motion studies suggest thatfaults in the mid-ocean ridge are dip–slip and exten-sional. The situation at trenches is more complex. Firstmotion studies of earthquakes along the Benioff Zone

show that faulting takes place at roughly 45 ° to theinclined surface of the downgoing slab. At depths ofgreater than about 300 km, the focal mechanisms arecompressional, but at shallower depths they are tensional.This pattern supports the view that in the lowermost partof the subducting slab, the plate experiences compressionas it is forced into a zone of greater viscosity or strengthat depth. The upper part, however, is in a state of tensionbecause of the gravitational body forces on the cool plate“hanging” from its upper edge. This force constitutesslab-pull. Isacks et al. (1969) provide a detailed analysisof the use of seismological studies at sites of subduction.

The relative motion of plates with constructive, con-servative, and destructive plate margins creates a contin-

ually changing picture of continental splitting, oceanbasin creation, ocean closure, and continental collision.This cycle of plate motion involving the birth and closureof oceans is termed the Wilson cycle since it is based onearly ideas of the opening and closing of the Atlanticocean by John Tuzo Wilson (Wilson 1966) (Fig. 1.8).Many sedimentary basins can be fitted into a particularphase of the Wilson cycle.

1.4 CLASSIFICATION SCHEMES OF SEDIMENTARY BASINS

Ideally, classifications are theories about the basis ofnatural order rather than dull catalogues compiled onlyto avoid chaos (Gould 1989, p.98, quoted in Ingersoll andBusby 1995, p.2). In this sense, classification schemes for sedimentary basins should both reveal something ofthe underlying mechanisms for basin development andreflect the natural variability of the real world.

Recent classification schemes of sedimentary basinsbased on plate tectonics have much in common. Theirlineage derives from Dickinson’s influential work in 1974which emphasized the position of the basin in relation tothe type of lithospheric substratum, the proximity of thebasin to a plate margin, and the type of plate boundarynearest to the basin (divergent, convergent, transform)(Fig. 1.9). The evolution of a basin could then be

NCKK, IB

M

P TNZ

CA, ALTNH NHKER IB NC

All others

100 km0

P

NC

TCA

NH

KK

ALT,ALK

M

IB

ALK0

100

200

300

Dep

th k

m

KERNZ

Fig. 1.7 Distribution of earthquake foci along Benioff zones, after Isacks and Barazangi (1977). NH, New Hebrides; CA, CentralAmerica; ALT, Aleutians; ALK, Alaska; M, Marianas; IB, Izu-Bonin; KER, Kermadec; NZ, New Zealand; T, Tonga; KK, Kuril-Kamchatka; NC, North Chile; P, Peru.

ABA1 11/22/04 4:31 PM Page 12

BASINS IN THEIR PLATE TECTONIC ENVIRONMENT 13

THE WILSON CYCLE

(a)

(b)

(c)

(f)

(e)

(d)Continental rift

New ocean crustin proto-oceanic

trough

Oceanic trench

Initiation ofsubduction

Closing of ocean basin

Subduction of ridge

Retroarc basinIntracontinentalstrike–slip faults

Fold–thrust beltand foreland basin

Continental collision andsubduction of continental lithosphere

Oceanic abyssal plain

Ocean basin withmid-ocean ridge

Passive continentalmargin

Stretchedcontinental

margin

Seafloorspreading

centre

Convergent margin withaccretionary prism

Forearc basins

Fig. 1.8 The Wilson cycle of ocean formation and ocean closure. Continental extension (a) is followed by the creation of a newoceanic spreading centre (b) and ocean enlargement (c). Subduction of ocean floor (d) leads to closure of the ocean basin. Subduc-tion of the oceanic ridge (e) takes place before continent–continent collision ( f ).

explained by changing plate settings and interactions.Dickinson (1974) recognized five major basin types onthis basis: (i) Oceanic basins, (ii) rifted continentalmargins, (iii) arc–trench systems, (iv) suture belts, and (v)intracontinental basins.

Strike–slip or transform related basins were conspicu-ously missing as a distinct basin type in this classification,a deficiency corrected in Reading (1982). Bally (1975)and Bally and Snelson (1980) differentiated three differ-

ent families of sedimentary basins based on their loca-tion in relation to megasutures, which in this context canbe defined to include all the products of orogenic andigneous activity associated with predominantly compres-sional deformation. The boundaries of megasutures areoften associated with subduction, whether it be of slabsof oceanic lithosphere (Benioff or B-type subduction) orof relatively buoyant continental lithosphere (Amferer orA-type subduction) and may also be the sites of impor-

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14 CHAPTER 1

tant wrench tectonism along transform faults. Ingersolland Busby (1995) developed the classifications ofDickinson (1974) and Ingersoll (1988) to recognize 26 different types grouped into classes of divergent settings,intraplate settings, convergent settings, transform set-tings, and hybrid settings.

Industry-based classifications are typified by thescheme suggested by Halbouty et al. (1970) and laterdeveloped by Fischer (1975) and Klemme (1980).Klemme’s scheme recognizes eight main types of basinbased on their architectural characteristics such as lin-earity, asymmetry, cross-sectional geometry, which arethemselves related to the tectonic setting and basin evo-

lution. The goal of categorizing a sedimentary basin andthereby gaining some predictive insights into frontierbasins is common to industry classifications such as those of Huff (1978) and Klemme (1980). It is pursuedby an Exxon group (Kingston et al. 1983a, b) to theextent of devising a formula for each basin, thereby facil-itating easy comparisons between basins and providingan “instant” idea of hydrocarbon potential. This classification system (Fig. 1.10) once again places basinsprimarily in their plate tectonic setting (lithospheric sub-strate, type of plate motion, and location on plate),reminiscent of Dickinson’s analysis over a decade earlier,and categorizes a basin according to three critical factors:

Intracratonic sagsand coldspot basins

Oceanic abyssal plains

Intracontinental pull-aparts

Continental rifts

Failed rifts(aulacogens)

Convergence DivergenceStrike –Slip or Neutral

Con

tinen

tal c

rust

Oce

anic

cru

stTr

ansi

tiona

l

Peripheral and retroarcforeland basins

Oceanic rifts

Backarc basins

Forearc basins

Thrust-sheet-topbasins (piggy-backbasins)

Intra-arc/intermontanebasins

Oceanic trench

Passive margins

Trench-slope basins

Plate margins

Plate interiors

Oceanic moats(surrounding seamounts)

Interplate continentaltransform-related basins

Interplate oceanictransform-related basins

TYPE OF PLATE MOTION

TYPE

OF

LITH

OSP

HER

IC S

UBS

TRAT

E

Fig. 1.9 Classification of basins using the type of lithospheric substrate, type of plate motion, and location with respect to the plateboundary.

ABA1 11/22/04 4:31 PM Page 14

Typ

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.

ABA1 11/22/04 4:31 PM Page 15

16 CHAPTER 1

the basin-forming tectonics, the depositional sequencesfilling the basin, and the basin-modifying tectonics.

1.4.1 Basin-forming mechanisms

Although the classifications outlined above undoubtedlyhave their uses, particularly in predicting source pres-ence, reservoir quality, availability of traps, etc., they havethe effect of scrambling some of the essential differencesand similarities between basins from the point of view oflithospheric mechanisms. Ingersoll and Busby (1995) rec-ognized six subsidence mechanisms, operating to differ-ent degree in their 26 basin types (Table 1.1), which canbe summarized as:• Crustal thinning, such as caused primarily by stretch-

ing or surface erosion;• lithospheric thickening, such as caused by cooling fol-

lowing stretching or accretion of melts derived fromthe asthenosphere;

• sedimentary and volcanic loading causing isostaticcompensation;

• tectonic (supracrustal) loading causing isostatic compensation;

• subcrustal loading caused by subcrustal dense loadssuch as obducted mantle flakes or crustal densificationdue to phase changes;

• asthenospheric flow primarily due to the subduction ofcold lithospheric slabs.

From the point of view of fundamental lithosphericprocesses, the major mechanisms for regional subsidenceand uplift (isostatic, flexural, and dynamic) can be summar-ized as follows (Fig. 1.11):

Isostatic consequences of changes in crustal and litho-spheric thickness; the thickness changes may be broughtabout purely thermally by cooling of lithosphere, forexample where new oceanic lithosphere moves awayfrom spreading centres. Thickness changes causing thin-

ning may be caused by mechanical stretching, subaerialerosion or at depth by delamination or the removal(stoping) of a deep lithospheric root. Mechanical thicken-

ing of crust and lithosphere, as in zones of continentalconvergence, generally causes isostatic uplift. Thickeningof the lithosphere by cooling, however, causes subsidence.

Loading and unloading at the surface and in the subsur-face, including the far-field effects of in-plane stresses;loading of the lithosphere may take place on a small scalein the form of volcanoes or seamount chains, and on a

large scale in the form of mountain belts, causing flexure

and therefore subsidence. The sediment infilling a basinalso acts as a sedimentary load, amplifying the primarydriving mechanism.

Dynamic effects of asthenospheric flow, mantle convec-tion, and plumes; subsidence or uplift are caused by the buoyancy effects of changes in temperature in themantle. Since these temperature changes are transmittedby viscous flow, the surface elevation changes may betermed dynamic.

For a given basin type, some or all of the mechanismsgiven above may have a major or minor role. Conse-quently, rather than attempting an encyclopedic cover-age of all basin types, we focus on the main lithosphericprocesses in Part 2 of this book. After an introductorychapter on the fundamentals of lithospheric mechanics(Chapter 2), basin forming processes are considered byinvestigating basins primarily caused by lithosphericstretching (Chapter 3) and basins primarily caused byflexure (Chapter 4). Chapter 5 discusses the role ofmantle–lithosphere interactions in basin development.Basins related to strike–slip deformation are consideredin Chapter 6.

Basins formed by stretching or thinning of the contin-ental lithosphere fall within an evolutionary sequence(Kinsman 1975; Veevers 1981). The early stages of thesequence correspond to the development of intracratonicrifts often associated with crustal doming. Such rifts mayevolve into oceanic spreading centres or may be abortedto form failed rifts or aulacogens. With seafloor creationand drifting of the continental edge away from thespreading centre, passive margin basins develop. Thesequence has been termed the rift-drift suite of sedimen-tary basins. The mechanisms of interest within this evo-lutionary sequence are therefore primarily the thermaland mechanical behavior of the lithosphere undertension, and the thermal contraction of the lithospherefollowing stretching.

Basins formed by flexure fall into two groups. Flexureof oceanic lithosphere as it approaches subduction zonesis responsible for the formation of deep oceanic trenches.It was the investigation of the deflection of the oceaniclithosphere at arc–trench boundaries that provided muchof the framework for the general theory of lithosphericflexure. Flexure of the continental lithosphere in contin-ental collision zones gives rise to foreland basins. Theforce system that causes flexure can be varied. In the caseof ocean trenches, it is probably a combination of gravi-tational body forces on the downgoing oceanic slab and

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Table 1.1 Basin classification adapted from Ingersoll and Busby (1995), modified from Dickinson (1974, 1976) and Ingersoll(1988), with modern and ancient examples.

Relative Basin type Basin description Modern Ancient exampleplate examplemovement

Divergent Terrestrial rift Rifts in continental crust, commonly Rio Grande Keeweenawansettings valleys with bimodal volcanism Baikal (Precambrian)

Rhine-Bresse Karoo (Jurassic)Graben Viking and Central

Grabens (Mesozoic)Proto-oceanic Incipient ocean basins floored by Red Sea East Greenland

rift troughs new oceanic crust, flanked by Gulf of California (Jurassic)young rifted continental margins

Intraplate Continental Mature rifted intraplate continental East coast, USA Early Paleozoic ofsettings rises and margins at continental–oceanic USA and Canadian

terraces boundary CordilleraContinental Progradation of sedimentary wedges Mississippi, Gulf Early Paleozoic

embankments at edge of rifted continental Coast, USA Meguma terrane,margins Canadian

Appalachians (?)Intracratonic Broad cratonic basins, commonly Chad Basin Paleozoic Michigan

basins with underlying rifts (Cenozoic) BasinCongo Basin Illinois Basin

Williston Basin (USA)Continental Stable cratons with thin, extensive Barents Sea Middle Paleozoic,

platforms sedimentary cover North AmericanMidcontinent

Active ocean Basins floored by oceanic crust at Pacific Ocean Various ophiolite-basins active divergent plate boundaries bearing complexes

(Semail, Oman),NeoproterozoicArabian Shield

Oceanic islands, Sedimentary aprons and platforms Emperor–Hawaii Mesozoic Snowaseismic ridges in intra-oceanic settings seamounts Mountain Volcanicand plateaus Complex

(Franciscan, California)

Dormant ocean Basins floored by oceanic crust, Gulf of Mexico Phanerozoic Tarimbasins neither spreading nor subducting Basin (China)

Convergent Trenches Deep troughs formed by subduction Chile Trench Cretaceous, Shumaginsettings of oceanic lithosphere Island (Alaska)

Trench–slope Structurally confined basins on Central America Cretaceous Cambriabasins subduction complexes Trench slab (California)

Forearc basins Basins within arc–trench gaps Sumatra Cretaceous GreatValley (California)

Intra-arc basins Basins along arc platform, including Lago de Early Jurassic, Sierrasuperimposed and overlapping Nicaragua Nevada (California)volcanoes

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Table 1.1 Continued

Relative Basin type Basin description Modern Ancient exampleplate examplemovement

Backarc basins Oceanic basins behind intra-oceanic Marianas Jurassic Josephinemagmatic arcs, and continental ophiolite (California)basins behind continental margin magmatic arcs without foreland fold–thrust belts

Retroarc Foreland basins on continental sides Andes foothills Cretaceous Sevierforeland basins of continental margin arc–trench foreland (Wyoming-

systems Utah)Remnant ocean Shrinking ocean basins between Bay of Bengal Pennsylvanian-

basins colliding continental margins Permian Ouachitaand/or arc–trench systems Basin(eventually subducted or deformed)

Peripheral Foreland basins superimposed on Persian Gulf Tertiary North Alpineforeland basins rifted continental margins during Indo-Gangetic Foreland Basin

continental collision Plain (Switzerland)Po Basin (Italy)

Piggy-back Basins carried above moving thrust Peshawar Basin Neogene, Apennines(thrust sheet sheets (Pakistan) (Italy)top) basins Meso-Hellenic Trough

(Greece)Foreland Basins formed among basement Sierras Laramide basins

intermontane cored uplifts in foreland settings Pampeanas (USA)basins (broken (Argentina)forelands)

Transform Transtensional Basins formed by local extension Salton Sea, Carboniferoussettings basins along strike–slip fault systems California Magdalen Basin (Gulf

of St Lawrence)Transpressional Basins formed by local compression Santa Barbara Miocene Ridge Basin

basins along strike–slip fault systems Basin (California)(California)

Transrotational Basins formed by rotation of crustal Western Aleutian Miocene Los Angelesbasins blocks about vertical axes within forearc Basin, California

strike–slip fault systems

Hybrid Intracontinental Basins on continental crust Qaidam Basin Pennsylvanian-settings wrench basins associated with strike–slip (China) Permian Taos Trough

tectonics caused by distant (New Mexico)collisional processes

Aulacogens Former failed rifts reactivated during Mississippi Paleozoic Anadarkoconvergent tectonics Embayment Basin (Oklahoma)

Impactogens Rift basins caused by stresses Baikal Rift Rhine Graben (Europe)transmitted from convergent plate (Siberia)margin

Successor Basins in intermontane settings Southern Basin Paleogene Sustut Basinbasins following cessation of orogenic and Range (British Columbia)

activity (Arizona)

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the excess in mass of the magmatic arc. In the case offoreland basins it is a combination of topographic loadsrepresented by the mountain belt, lateral density varia-tions of lithospheric material (caused, for example, byprior stretching of the continental lithosphere in theoverridden plate or by obduction of dense mantle flakes)and horizontal “in-plane” forces set up by the shallow“end-on” collision of buoyant continental lithosphere.

Processes within the mantle have an important role inbasin development. Mantle plumes originating at thecore–mantle boundary impinge on the base of the litho-sphere and spread out laterally over a length scale of103 km. They are instrumental in continental splittingand the formation of new ocean basins, and have a major

role in igneous underplating and isostatic regional uplift.Present day hotspots over plumes or ascending limbs ofmantle convection systems are characterized by topo-graphic doming and commonly rifting. Basins locatedover downward limbs of convection systems, cold-spot

basins, appear to be broad, gentle sags. The onset ofsubduction of cold oceanic slabs at ocean–continentboundaries causes far-field tilting of the continental platetowards its margin and therefore has a potentially majoreffect in forearc, intra-arc, and retroarc settings. In addi-tion, it is believed that supercontinental assemblies in thegeological past have experienced very long wavelengthtopographic doming caused by elevated sublithospherictemperatures generated by an overlying insulating lid.

BASINS IN THEIR PLATE TECTONIC ENVIRONMENT 19

Foreland basins

Rifts

Postrift andpassivemargins

Flexural

Lithosphericthicknesschanges(stretching)

Lithospheric cooling

Dynamic (mantle)

Supracrustal, sub-crustal, end loads,and torques

Dynamic topography

Isostatic response to crustal thinning;brittle fault displacements

Induced convection

Thermal contraction/re-equilibration of lithosphere

Sediment loads

Isostatic

Crustal flakes, push-up ranges

Mantle involvementin local stretching

Pull-aparts inPrincipal

Displacement Zone

Strike–slip basins

Fig. 1.11 Fundamental mechanisms of basin formation: flexural, isostatic, and dynamic. The importance of these mechanisms inforeland, rift, postrift and passive margin, and strike–slip basins is indicated by the size of the boxes.

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SUMMARY

A knowledge of the behavior of the lithosphere is essen-tial if we are to understand the initiation and develop-ment of sedimentary basins. Lithospheric processes are responsible for the highly dynamic nature of the tectonic processes near the surface of the Earth. TheEarth’s outer layer or lithosphere can be regarded as athermal boundary layer between the cool atmosphere oroceans and the hot interior. The lithosphere is thereforea thermal entity but it also has a physical significance.The upper half of the lithosphere is sufficiently rigid that it is able to store and transmit stresses. Because the deformations caused by applied forces are generallyrecoverable, this outer zone is known as the elastic lithosphere.

Applied forces of whatever origin cause stresses whichresult in deformation or strain. The simplest view ofapplied forces is of those due simply to the weight of anoverlying rock column, known as lithostatic stress. The difference between the actual stress and the lithostaticstress is a tectonic contribution known as deviatoric stress.Deviatoric stresses can be either tensile or compressive.

In an elastic solid there is a clear relationship betweenthe stresses and resultant strains. The exact relationshipdepends on material properties known as Young’smodulus and Poisson’s ratio. Where only one of the prin-cipal axes is nonzero, a uniaxial state of stress is said tooccur and the relation between stress and strain is calledHooke’s law. If there are two nonzero components ofprincipal stress, we have the condition termed plane stress.In an analogous fashion, uniaxial strain and plane strain

refer to the coordinate system of principal strains. In astate of stress where all the principal stresses are equal,an isotropic state of stress, the fractional volume change

caused by isotropic compression is given by the bulk

modulus or its reciprocal, the compressibility.The lithosphere is able to bend. The shape of the

flexure of the lithosphere depends on its rigidity and thenature of the applied force or load causing the bending.Bending is accompanied by longitudinal stresses andbending moments in the plate. The bending moment isrelated to the local radius of curvature by a coefficientcalled flexural rigidity.

Whereas the upper 15–50km of the Earth is thoughtto behave more or less elastically over geological timescales, rocks below this level deform with a creep-likebehavior and thus relax stresses.

In order to understand the mechanical behavior ofthe Earth it is necessary to understand its thermal struc-ture, since this determines its rheology. In the lithosphere,heat transfer is predominantly by conduction, whereas inthe mantle convection is extremely important. Convec-tion is the natural result of differential heating above a critical Rayleigh number and estimates for Earth’smantle strongly suggest that convection must be takingplace. A thermal boundary layer develops along theupper surface of the convecting fluid because heat is lostby conduction to the surface. This cool boundary layer,the lithosphere, detaches and sinks due to gravitationalinstability along subduction zones. For conduction, theheat flux is related to the temperature gradient by a coefficient, the thermal conductivity. Heat fluxes in the continents are determined primarily by conduction from radioisotopic heat sources, whereas in the oceansthey reflect cooling of newly created oceanic lithosphere.The variation of temperature with depth is known as the geotherm. Measurements of continental heat flowsindicate a linear relationship between radiogenic heatgeneration and surface heat flow, with the intercept

C H A P T E R

2 The physical state of the lithosphere

The kindly Earth shall slumber, lapt in universal law

(ALFRED LORD TENNYSON, IN MEMORIAM (1850))

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