Principles of Stratigraphy

Principles of Stratigraphy Michael E. Brookﬁeld Principles of Stratigraphy Principles of Stratigraphy Michael E. Bro...

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Principles of Stratigraphy Michael E. Brookﬁeld

Principles of Stratigraphy

Principles of Stratigraphy Michael E. Brookﬁeld

© 2004 by Blackwell Publishing Ltd 350 Main Street, Malden, MA 02148-5020, USA 108 Cowley Road, Oxford OX4 1JF, UK 550 Swanston Street, Carlton, Victoria 3053, Australia The right of Michael E. Brookﬁeld to be identiﬁed as the Author of this Work has been asserted in accordance with the UK Copyright, Designs, and Patents Act 1988. All rights reserved. No part of this publication may be reproduced, stored in a retrieval system, or transmitted, in any form or by any means, electronic, mechanical, photocopying, recording or otherwise, except as permitted by the UK Copyright, Designs, and Patents Act 1988, without the prior permission of the publisher. First published 2004 by Blackwell Publishing Ltd Library of Congress Cataloging-in-Publication Data Brookﬁeld, M. E. (Michael E.) Principles of stratigraphy / Michael E. Brookﬁeld. p. cm. Includes bibliographical references and index. ISBN 1-4051-1164-X (pbk. : alk. paper) 1. Geology, Stratigraphic. I. Title. QE651.B787 2004 551.7–dc21 2003002573 A catalogue record for this title is available from the British Library. Set in 9/11pt Photina by SNP Best-set Typesetter Ltd., Hong Kong Printed and bound in the United Kingdom by TJ International Ltd, Padstow, Cornwall For further information on Blackwell Publishing, visit our website: http://www.blackwellpublishing.com

Contents

Preface Acknowledgments 1

Introduction 1.1 Stratigraphy – why bother? 1.2 Development of stratigraphy 1.3 Phases of study

Part I Basics

viii x 1 1 3 8

11

2

Weathering 2.1 Types of weathering 2.2 Rates of weathering 2.3 Soil formation 2.4 Weathering and soil formation under water

13 13 15 16 21

3

Sediments and sedimentary rocks 3.1 Transportation and deposition 3.2 Clastic sediments and sedimentary rocks 3.3 Chemical and biochemical sediments and sedimentary rocks

22 22 24

4

Major environmental complexes and their recognition 4.1 Introduction 4.2 Impact and volcanic environments 4.3 Continental environments 4.4 Environments under water 4.5 Mixed environments 4.6 Peculiar environments

46

66 66 67 67 81 99 99

vi Contents

Part II Tracing environments in space and time

101

5

The vertical dimension 5.1 The local section 5.2 Breaks in the record 5.3 Dividing the local section: the type section 5.4 Strata and stratiﬁcation

103 103 105 111 114

6

The horizontal dimension 6.1 Physical correlation 6.2 Lateral changes 6.3 Mapping

115 115 127 134

7

The time dimension 7.1 Age equivalence 7.2 Relative ages 7.3 Numerical methods (ages in years) 7.4 Calibration of relative and numerical dates

140 142 142 158 170

8

Basin analysis 8.1 Basin-ﬁll architecture 8.2 Sediment provenance 8.3 Paleocurrents and sediment dispersal 8.4 Backstripping 8.5 Paleothermometry 8.6 Paleogeographic and paleotectonic maps

171 171 175 175 175 180 184

9

Stratigraphic systems 9.1 Development of the stratigraphic system 9.2 Cycle stratigraphy 9.3 Genetic sequence stratigraphy 9.4 The current system

186 186 190 192 197

Part III Interpreting geologic history

207

10 Tectonics 10.1 Geodesy 10.2 Hypsometry 10.3 Gravity 10.4 Isostasy 10.5 Tectonics and sedimentary basins 10.6 Exotic terranes 10.7 Terrane analysis of orogenic belts

209 209 210 210 213 218 235 235

11 Sea-level changes 11.1 Eustatic or “absolute” changes of sea level 11.2 Relative changes of sea level

241 241 246

Contents vii 12 Climate 12.1 Present distribution and character of climates 12.2 Identiﬁable climatic effects on sediments, biotas, and stable isotopes 12.3 Controls on climate and climatic change

255 255

13 Biology 13.1 Atmosphere and ocean changes 13.2 Bioclastic sediment changes through time 13.3 Sediment mixing by organisms 13.4 Biogeographic changes

271 271 271 272 272

14 Stratigraphic problem times and places 14.1 Quaternary stratigraphy 14.2 Archeological stratigraphy 14.3 Proterozoic stratigraphy 14.4 Archean stratigraphy

276 276 281 283 284

15 Extraterrestrial stratigraphy 15.1 The Moon 15.2 Mercury 15.3 Mars 15.4 Venus

285 285 287 287 287

Appendix 1 Imperial/metric conversions Appendix 2 Figure legends Appendix 3 Geologic time-scale

292 293 296

Glossary References Index

297 307 321

Color plates fall between pp. 182 and 183.

259 265

Preface “Stratigraphy may be deﬁned as the complete triumph of terminology over facts and common sense.” P.D. Krynine

Stratigraphy is one of the most demanding, fundamental, and interesting of geologic disciplines. It tries to reconstruct history with a few basic principles, requires both careful observation and wide-ranging imagination, and involves many individually fascinating subdisciplines. Yet, as the quotation above shows, this is not how it appears to many people. Stratigraphy has a poor image because it is sometimes presented only as a way of classifying and organizing strata with “codes.” However, its basic principles are very simple and applicable in any study in which a history must be reconstructed from layers. Thus, stratigraphic principles are an essential part of archeology and even forensic science (the sequence of events during a murder, for example). They are now used to work out the history of planetary surfaces in studies that relive the early 19th century excitement in exploratory mapping and correlation (Carr et al. 1993). This book is intended to back up a second-year university course, so I assume that students using it have already taken a basic Introductory Geology course (or its equivalent), though some material is repeated for convenience (since students tend to forget things over time). I wish to emphasize the need to: (i) know how stratigraphy developed and how this constrains current approaches and dogmas; (ii) study all aspects of modern processes and environments (including extraterrestrial ones) as a basis for paleoenvironmental interpretation; (iii) understand the principles of correlation and dating as a basis for stratigraphic reconstructions; (iv) know

some basics of geophysics, tectonics, climatology, and paleobiology in order to explain the development of sedimentary basins; and (v) make ﬁeld observations in person, to understand the limitations of data and interpretations. First, it is essential to know the history of ideas in a subject, and how and why choices were made among different alternatives. The current controversies of sequence stratigraphy simply recapitulate the disputes between Werner and Hutton in the 18th century, between Oppel and Gressly in the 19th century, and between Grabau and Ulrich in the early 20th century. Basic stratigraphic concepts were only slowly established after much discussion. Perhaps the most important thing to understand is why discredited ideas were accepted at the time (and vice versa), as this might hopefully lead to much needed humility and openmindedness in the face of current dogma (Menard 1986; Raup 1986). Second, the conditions under which ancient stratiﬁed rocks formed can only be worked out by studying their modern counterparts (where possible) and deriving principles from sedimentology and ecology that might help in this (Fraser 1989). Third, the relationships of environments in space and time require speciﬁc concepts of correlation and dating which are peculiar to stratigraphy (Berry 1987). Fourth, it is impossible to interpret the stratigraphic development of an area, or sedimentary basin, without knowing some basic concepts in geophysics, tectonics, climatology, and biology.

Preface ix Fifth, personal observations on actual rocks are essential, otherwise people get completely out of touch with their material. Arkell (1933, p. 36) cites the example of the famous Jurassic geologist, S.S. Buckman: “That Buckman, who had tramped the Cotswolds and the Sherborne country from end to end and knew every quarry intimately, whose earlier work was built up solely on sound ﬁeld-work, could also be the author of his last paper . . . and some of the later parts . . . of Type Ammonites is difﬁcult to believe. Without any practical knowledge of the Cornbrash, without describing so much as a single section, he proceeded to divide it up into 11 brachiopod zones and coined for it 5 new stage names. Neither zones nor stages have any foundation in fact.” There has always been disagreement on how to teach stratigraphy, and this reﬂects the fundamental disagreements of earlier geologists. Werner’s historical view of earth history, marked by discrete events in time with a starting point and an end point, contrasted with Hutton’s view of unending and unchanging cyclical processes; time’s arrow versus time’s cycle (Gould 1987). Though both traditions are found in modern textbooks, one or other tends to dominate and each author’s experience naturally colors their individual approaches. The classiﬁcation and correlation tradition is found in Lemon (1990), Prothero (1990), and Schoch (1989). The environments through time approach is found in Blatt et al. (1991), Boggs (2001), Brenner and

McHargue (1988), Hallam (1981), Matthews (1984), Nicholls (1999), and Prothero and Schwab (1996). I am mostly sympathetic with the environments through time approach; yet this sometimes ignores the essential (perhaps boring) need to actually correlate and date rocks, and understand the problems in doing that. It also tends to minimize irreversible, sometimes rapid and catastrophic changes. Thus, certain lithologies are conﬁned to speciﬁc times (Ager 1973): chalks are found only from Cretaceous times onwards (after the evolution of carbonate-secreting plankton). And, as Alvarez et al. (1980) established, large meteorite impacts have often affected the stratigraphic record. I wrote this book hoping to instil in students the usefulness, wonder, and relevance of stratigraphy. I also wrote it to re-establish the necessity for tectonic and geophysical knowledge to interpret the stratigraphic record. It is truly amazing that some current stratigraphy texts have nothing on isostasy: a text such as Whole Earth Geophysics (Lillie 1999) should be compulsory reading for all stratigraphers. I have tried to give references for all main ideas and also for the ﬁgures. Students should not accept statements in which sources are not given. It is extremely important that students get into the habit of questioning and checking both facts and ideas, especially from their teachers. Otherwise, unquestioned, authority-derived dogma tends to dominate a subject – remember the history of the continental drift theory (Oreskes 1999).

Acknowledgments

All books are the result of an author’s experience. I learned most from three rather different teachers – D. Ager, J.R.L. Allen, and A. Hallam – all of whom, however, emphasized the derivation of principles from carefully chosen observations, rather than the simple accumulation of data, or the mindless application of formulae. I am grateful to several generations of students in my stratigraphy classes who pointed out ways of simplifying and clarifying the text: there is not much point in producing a book that students neither understand nor read. I also owe a debt to all authors of other stratigraphy texts, whose ideas and methods have helped in writing my own, and thanks to a variety of anonymous

readers who improved the manuscript with their comments. Of the individuals who helped me during the writing, I must ﬁrst thank my wife, Kathleen, for support over many years and for contributing to writing and editing during the early stages. Steve Sadura and Don Irvine at Guelph did wonderful jobs in producing respectively the glossary and artwork. The staff at Blackwells, speciﬁcally Ian Francis, Delia Sandford, Rosie Hayden, Cee Pike, Linda Auld, and Sue Worrall, managed to be both prompt and efﬁcient (something I am incapable of) in transforming a rough manuscript into a ﬁnished book.

1 Introduction

1.1 Stratigraphy – why bother? 1.2 Development of stratigraphy 1.3 Phases of study

1.1 Stratigraphy – why bother? Stratigraphy gives you techniques for working out earth history: it integrates diverse materials into a coherent view of how the earth and its life forms evolved. Though stratigraphy (literally writing about strata) is mostly about working out the history of sedimentary rocks, in order to do this you also need to know the effects of magmatism, metamorphism, tectonism, climatic change, and sea-level changes, and the effects of organic evolution. So, stratigraphy integrates data and concepts from many specialties, and in practice ends up as a much more comprehensive study than its name implies. Stratigraphy also lets you test ideas on how varying combinations of processes affect the planets through time. For example, as evidence for continental drift and changing climates, Wegener (1915) used the presently separated positions of Carboniferous Mesosaurusbearing and glacial sediments, which were most plausibly explained by an originally compact supercontinent. Together, history and process let you work out how, when, and why environments changed through time.

Stratigraphy, perhaps most importantly, also helps you to understand how many economic materials formed and got distributed in the way they did – and so will hopefully help you ﬁnd more. For example, the Wembley Field is one of many isolated oil and gas reservoirs in the Middle Triassic in Alberta, Canada (Fig. 1.1). Finding out why the oil is there, and where other similar oil and gas ﬁelds are, requires you to proceed logically through the various phases of stratigraphy. First, what are the actual oil-bearing rocks and how were they deposited? The local (borehole) sections mostly consist of porous sands alternating with clays arranged in coarsening-upwards cycles, deposited by waves and currents as marine barrier island deposits (Fig. 1.2). Second, how are these sediments arranged spatially and how old are they? In the absence of outcrops, space correlations have to be worked out from borehole logs and seismic sections which show the arrangement and thickness of the strata and environments and that the productive oil and gas wells are in linear sand bodies of a particular type of marine barrier bar (Fig. 1.3).

Fig. 1.1 Location of Middle Triassic oil and gas ﬁelds in west-central Alberta, Canada (modiﬁed from Willis & Moslow 1994a, ﬁg. 3). (AAPG © 1994. Reprinted by permission of the AAPG whose permission is required for further use.)

Fig. 1.2 Section of Halfway Formation with barrier-bar interpretation, and porosity and permeability of the oil- and gas-bearing units (modiﬁed from Willis & Moslow 1994a, ﬁg. 5). (AAPG © 1994. Reprinted by permission of the AAPG whose permission is required for further use.)

Introduction 3

Fig. 1.3 Sand thickness map of Halfway Formation (from Willis & Moslow 1994a, ﬁg. 6). (AAPG © 1994. Reprinted by permission of the AAPG whose permission is required for further use.)

Further detailed studies showed that the oil and gas is concentrated speciﬁcally in the well-sorted and porous sandstones deposited in active tidal inlets and ebb-tidal deltas whose distribution could be plotted in sections (Fig. 1.4) and on thickness maps (Fig. 1.5). But why is the oil and gas concentrated in the Halfway Formation, when many other underlying units also show coarsening-upwards barrier-bar sections? The reason is the peculiar development of this unit. Small rises in sea level, which led to the formation of transgressive barriers from eroding shoreline sand deposits, alternated with periods of stable sea level, during which seaward migration of barriers led to burial of the transgressive barriers with ﬁne-grained impermeable backbarrier and coastal plain sediments (Fig. 1.6). The oil and gas reservoirs can be located by tracing the condensed radioactive clays at the base of each marine transgression up-dip into the transgressive barriers. Third, what is the history of the area? Regional stratigraphic studies show that it formed part of a subsiding passive margin shelf in western Canada (Fig. 1.7). Inherently unpredictable evolution of discontinuous barrier islands was controlled by the interactions of tides and sea-level changes. The ages come from marine fossils ﬁtted into the standard geologic time-scale. Fourth, how does this area ﬁt into the overall interpretation of earth history? Although working this out is not always necessary in a regional study, the results will contribute to understanding the overall development

and character of the Triassic. This study of the Wembley Field used essentially a modern physical stratigraphy approach to the distribution of strata and their environmental interpretation. Biological, climatic, and tectonic factors were not used (and were probably not needed) to understand the ﬁeld. Nevertheless, all possible factors and approaches should be considered both before and during any study, since any individual is usually biased towards those approaches made familiar and comfortable by education and experience. On the one hand, a stratigrapher should know why he works in a particular way and realize that other ways may be equally valid: many controversies arise because people do not appreciate the aims of different stratigraphic studies. On the other hand, a stratigrapher should criticize inappropriate and confusing methods and concepts traditionally applied to his area or period: much dead weight and confusion could be lifted by applying alternative methods and concepts. Furthermore, it is impossible to understand the traditions followed in different areas and periods and evaluate their results without knowing how stratigraphy developed (Gohau 1991).

1.2 Development of stratigraphy Geology became a speciﬁc discipline at the end of the 18th century with the description of strata as its focus,

4 Chapter 1

Fig. 1.4 Dip section across Wembley Field showing tidal channels, ebb-tidal delta environments, and oil- and gas-producing intervals (from Willis & Moslow 1994b, ﬁg. 13).

Fig. 1.5 Distribution of tidal inlet sand units (from Willis & Moslow 1994b, ﬁg. 15).

and developed from the Romantic movement’s taste for savage nature and travels to remote places (as it still does: I work in the mountains of Central Asia). Stratigraphy, like keeping a journal or collecting beetles, gave serious purpose to tours that might otherwise seem aim-

less or frivolous (Porter 1977). The ﬁrst thing to do was work out the superposition of strata in an area. Such local successions, securely based in limited areas, then served as “types” for similar successions elsewhere. Scientiﬁc order could then be imposed worldwide on the basis of rock and fossil similarities. Most people spent their time tracing and mapping strata, and in trying to classify the geologic column rather than trying to explain it (Secord 1986). This is the ﬁrst tradition of classiﬁcation and nomenclature. Within this tradition, the emphasis on the rock strata and the fossils they contained varied greatly and led to different stratigraphic approaches or styles. Some people emphasized the actual rocks (LITHOSTRATIGRAPHY), some the fossils the rocks contained (BIOSTRATIGRAPHY). Though the actual strata determined the aproach, personal style determined the actual ﬁeld area chosen for study. Scientiﬁc style was shown in everyday ﬁeld practice and not in grand generalizations, though these evolved together. And after formative undergraduate study and learning, the rocks studied and the styles chosen tended to narrow and ossify. The contrast between rock (lithostratigraphic) and fossil (biostratigraphic) approaches is best exempliﬁed in

Introduction 5

Fig. 1.6 (a–c) Model for the development of a transgressive barrier, and (d) location of successive transgressive barriers (from Willis & Moslow 1994a, ﬁgs 12, 13). (AAPG © 1994. Reprinted by permission of the AAPG whose permission is required for further use.)

the work of Sedgwick and Murchison, which led directly to the controversy over the Cambrian–Silurian boundary in the mid-19th century (Secord 1986). Sedgwick’s fame rested on his insight into structure, his ability to visualize rocks in three dimensions and interpret their relationships after only a few traverses: he emphasized the distribution of rock types, worked on the basically unfossiliferous strata of northwestern Wales, and hence used a lithostratigraphic approach. Murchison’s fame rested on his development of geologic systems based on fossils: he emphasized the vertical succession of faunas, named the Silurian and Permian periods, worked on the fossiliferous rocks of the Welsh–English border area, and used a biostratigraphic approach. What might have happened if they had changed places? It would be pointless to ask. They chose their work areas precisely because of their different approaches and the type of rocks present. Sedgwick worked in North Wales because its unfossiliferous schists and slates were structurally complex, and contrasting well-exposed rock types could be used to trace the structure. Murchison worked in the Welsh Borders because it was there that biostratigraphic divisions could be recognized and correlated in fossiliferous and structurally simple, but poorly exposed, and lithologically repetitive successions. In Europe, the biostratigraphic approach eventually overshadowed the lithostratigraphic approach because

of the success of d’Orbigny and Oppel in developing biostratigraphy and of Gressly in developing the FACIES concept. Both concepts evolved together and were based on the fortunate pecularities of European Jurassic rocks. The European Jurassic contains widespread, rapidly evolving, and easily recognizable lineages of swimming coiled cephalopods called AMMONITES. These are common in many different rock types, and biostratigraphic time units based on the vertical ranges of different ammonite species were erected independent of lithology by d’Orbigny (1842) and particularly by Oppel (1856–8). These “time zones” were often traceable across western Europe, even where the rock types changed, because the area is small enough to limit biogeographic effects on free-living organisms. Oppel was professionally perceptive (or lucky) in both the fossil group and the area he studied; he was personally less fortunate in dying of typhoid at the age of 35. Also, the European Jurassic was deposited during a period of continental rifting and splitting. The changes in rock types from one area to another reﬂected great lateral changes in environments, both vertically and horizontally, emphasized by compressional shortening in the type area of the Jura mountains. Simple tracing of rock units is difﬁcult or impossible because of the great lateral variation within even the small sizes of European countries. Thus, geologic surveys recognized from the

Fig. 1.7 Thickness (in meters) of Middle Triassic sediments on western Canadian margin (from Willis & Moslow 1994a, ﬁg. 1). (AAPG © 1994. Reprinted by permission of the AAPG whose permission is required for further use.)

Introduction 7 precise biostratigraphy available that some rocks were of the same age even when they were markedly different in both sediment type and benthic fossils. Gressly (1838–41), working in the Jura mountains of northeastern France, called these facies changes. The need to interpret Gressly’s sedimentary and biological facies changes led to the second tradition of using the sedimentology and ecology of modern environments to interpret ancient strata and fossils. Based on Lyell’s strict PRINCIPLE OF UNIFORMITARIANISM, this fundamentally modern approach to stratigraphy was established in Europe by the end of the 19th century and is best exempliﬁed in the work of J. Walther (1893–4). No lucky circumstances comparable with the European Jurassic existed in North America. Both facies and biostratigraphic zonation was practically ignored until the mid-20th century. In North America, the eastern seaboard was the ﬁrst to be studied geologically. Between 1832 and 1851, James Hall, the outstanding American geologist of his generation, moved westwards, describing sections and fossils from the Paleozoic as he went (Dott 1985). Hall’s studies proceeded in the same way as Murchison’s, except that lateral changes in lithology were nowhere near as obvious, practically all the fossils were benthonic and facies (environmentally) controlled, and no good fossil zonation was possible outside the graptolite-bearing shales of the Appalachians. The uniformity of the extensive fossiliferous sedimentary units almost forced the view that widespread lithologies with characteristic fossils succeed one another; and that distinct units bearing different fossils imply time differences. This view was forcefully promoted by Hall’s successor, E.O. Ulrich. Starting in 1885, Ulrich described the Lower Paleozoic of central and eastern North America; what we now call the craton interior and passive margin. According to Ulrich, the Paleozoic shelf seas occupied small, shallow, and often disconnected basins. These basins changed their extent and character depending on local rhythmic deformations with consequent transgressions and regressions. Each individual advance of the sea laid down a rock unit with relatively constant lithology and fauna. Each unit was terminated by nondeposition rather than by lateral change into a different lithology. Each unit was separated by a widespread time break marking retreat of the sea. These breaks would be commonplace yet inconspicuous because of the low relief of the continental interior, the shallowness of the shelf and interior seas, and the frequency of oscillations. Ulrich interpreted broad contemporary lateral changes

of sediment and fauna, reﬂecting lateral changes of environment, as distinct sedimentary lenses of different ages. This led to the concept of troughs and barriers, according to which deposits of sandstone, shale, limestone, etc., which formed an orderly contemporary facies change, were considered by Ulrich to have been deposited separately and successively in one or another of four or ﬁve parallel troughs (Merk 1985). On this basis, in 1911, Ulrich proposed two new systems, the Ozarkian and the Canadian, between the Cambrian and Ordovician systems. Ulrich’s ideas have since been castigated as prime examples of reactionary dogmatism that retarded the acceptance of the facies concept in America (Dunbar & Rogers 1957). True enough: but for some of the Paleozoic, Ulrich was right. For example, the shallow-water Ordovician of eastern North America does show separate and successive overlapping lenses separated by time breaks in some areas, albeit with more facies changes than Ulrich would accept (Brookﬁeld & Brett 1988). And Ulrich’s ideas have recently been resurrected in the concept of sequence stratigraphy. Ulrich’s friend, W.A. Grabau, promoted the opposite facies view. However, Ulrich’s ideas dominated American stratigraphy until the mid-20th century, and there is a residual tendency to downplay facies at the expense of layer-cake and cyclical stratigraphy even now; witness the success of cyclical and sequence stratigraphy, “punctuated aggradational cycles,” and “ecostratigraphy.” (Ulrich and Hall both got a separate chapter in the 1985 Geological Society of America Centennial Volume on the history of North American geology; Grabau is not even mentioned.) Grabau worked on the Silurian and Devonian of western New York State, on rocks later used as classic examples of facies change (see Chapter 6). However, until the early 20th century, rocks could still only be indirectly dated. To get a date a known accumulation or loss has to be divided by a known and uniform rate. So, people estimated the amount of salt now in the sea, divided it by the rate of supply, and got values of about 100 million years for the age of the oceans. Or they estimated the thickness of sediment preserved, divided it by the average rate of supply, and got values of about 150 million years since sediment started accumulating on the earth. All these efforts foundered on the unreliability of both accumulation values and uniformity of rate, and the undoubted removal of both salt and sediment. Estimates based on a molten cooling earth could not be faulted on 19th century physics, and Lord Kelvin’s ﬁnal 1897 estimate of 27 million years for the

8 Chapter 1

Fig. 1.8 Dynamic stratigraphy (after Aigner 1985, ﬁg. 1).

age of the earth was generally accepted. But in 1896 the discovery of radioactivity by Becquerel gave an additional source of heat, and demolished the basis for Kelvin’s short estimate. It also gave, for the ﬁrst time, a reasonably accurate way of dating rocks in years. With this discovery, the basic stratigraphic trilogy of rock type distribution, relative time, and absolute time was established.

1.3 Phases of study The phases of study in stratigraphy essentially follow the original development of stratigraphy in the 19th century and are as relevant now as when they ﬁrst appeared. These phases are followed in this book.

Phase 1: Basics The basics, needed before starting any stratigraphic studies, involve ﬁrst being able to: 1 identify and classify minerals, rocks, and fossils accurately; 2 infer the processes that formed the minerals, rocks, and fossils from ﬁeld and laboratory studies of the effects of modern physical, chemical, and biological processes; 3 recognize the ancient depositional (and rarely nondepositional) environments, by comparing the variety, intensity, and periodicity of processes in modern environments with those inferred in ancient rocks successions; 4 map the obtained data on maps and sections of various types.

Introduction 9 These basics have to be done well as they form the foundation for all further studies. After mastering these you can then go on to phase 2. Phase 2: Tracing environments in space and time Tracing environments in space and time requires four steps: 1 An overview of the area studied involves a preliminary survey of what work has already been done, together with an analysis of the type and distribution of the rocks and the problems in studying them. The ﬁrst can be done in a library and/or by talking with previous workers. The second involves areal studies of the surface using remote sensing (e.g. satellite and aerial photographs, although personally I like paragliding), and by actual reconnaissance on foot or by some form of transport; and studies of the subsurface with geophysics (e.g. seismic proﬁles) and boreholes. During this work, any problems of access and exposure should become apparent. This step overlaps and can help in planning the second step. 2 The description of local sections involves measuring STRATA, describing their attributes (including composition, texture, structure, and fossil content), and working out the processes that formed the sediments and the succession of depositional environments present in the sections. For this you need to know the basic principles of sedimentology and ecology. 3 The correlation of local sections in space and time involves the physical tracing and mapping of rock units, their relative dating by means of fossils and other methods, and dating by means of RADIOMETRIC DATING or cross-correlation with standard time-scales. You need to know how environments and organisms vary and differ

in a variety of sedimentary basins; how to plot these variations on maps and diagrams (and the advantages and limitations of different methods); how relative time units are constructed; and how radiometric and other methods of dating are done. 4 The reconstruction of sedimentary basin history involves synthesis, often on maps and crosssections, of trends in rock type, petrology, facies, thickness, and so on. The three steps following the initial overview are marvellously summarized by Aigner (1985) as stratinomic analysis (giving the depositional dynamics of the sediments), facies analysis (showing the lateral and vertical variability or facies dynamics of the sediments), and BASIN ANALYSIS (explaining how the basin evolved during sedimentation) (Fig. 1.8). Phase 3: Interpreting geologic history Interpreting geologic history involves evaluating the effects of controlling processes such as tectonics, sealevel changes, climate, and biology (the effect of organisms) on sedimentary basin history. It requires the ability to synthesize large amounts of data from many ﬁelds. On the grand scale it involves the correlation of histories of individual basins and intervening areas to give a worldwide picture of the development of a planet, and requires wider consideration of the stratigraphic peculiarities and problems of different time periods on these planets. Each of these phases depends on the competence of the preceding phase. Thus, poorly described local sections inevitably result in poor stratigraphy, poor environmental reconstructions, poor correlations, and lousy reconstructions.

I Basics “Only knowledge of facies relationships drawn from study of modern environments can save us from the barren cataloging of rock and fossil sequences that sometimes passes for stratigraphy.” Middleton (1973)

Examining Ordovician quartzites, Jebel Uweinat, NW Sudan.

The ﬁrst phase of stratigraphy, the study of actual strata, and the processes that form, transport, deposit, and modify strata, is an attempt to infer how ancient strata formed. First, you must understand how physical, chemical, and biological processes produce the variety of sediments found in modern environments. This requires a broad knowledge of sedimentology. Second, you must be able to describe and identify sedimentary rocks accurately with criteria useful in working out what processes formed them and how they have been modiﬁed after deposition. This requires knowledge of various systems of classiﬁcation and their basis, together with their advantages and defects. Third, you must be able to recognize major environmental complexes from the rather limited clues still left in ancient rocks. This requires a knowledge of how various processes combine to deﬁne a speciﬁc environment. Only then can you start deﬁning and mapping stratigraphic units. Unfortunately, only too often, description and mapping are done ﬁrst, before the student has much idea about how the sediments formed.

The ﬁrst part of this book is thus a necessary outline of the physical and biological processes that form, transport, and deposit sediments, and the rock and fossil evidence required to reconstruct them. The properties measured in sediments and sedimentary rocks should be those most useful in determining how they formed. You must understand these processes, and the features they produce, in order to realize the basis for classiﬁcation and to be able to recognize environments. For this reason, the discussion of the way sediments form precedes their classiﬁcation. Good introductions to the physical and biological processes that form soils and sediments are Fitzpatrick (1980), Knapp (1979), Taylor and Eggleton (2001), and Weyman and Weyman (1977). This book uses the metric system. Appendix 1 is a metric/imperial conversion table. Appendix 2 is a legend for symbols used in most ﬁgures. Appendix 3 is a geologic time-scale.

2 Weathering

2.1 2.2 2.3 2.4

Types of weathering Rates of weathering Soil formation Weathering and soil formation under water

Weathering (alteration in place) determines the nature of the sediment at the source. It begins with the physical and chemical breakdown of materials exposed at or near the surface (both on land and under water), and continues during transportation, deposition, and the migration of pore ﬂuids. At the surface, PHYSICAL WEATHERING and CHEMICAL WEATHERING proceed at different rates depending on rock type, climate, and slope.

2.1 Types of weathering Physical weathering is simply the breaking down of material into smaller pieces. Softer rocks and minerals can be worn into small pieces faster than harder minerals, while fractured rocks and minerals with cleavages break into smaller pieces faster than massive rocks and minerals. Talc (a soft sheet silicate with good cleavages) rapidly breaks down into powder; quartz (a hard framework silicate with no cleavage) survives to form the main mineral of sands.

An important consequence of physical weathering is that it increases the surface area available for chemical breakdown. Chemical weathering is more complicated. Agents such as water, carbon dioxide, oxygen hydrogen ions, bacteria, humic acids, etc., break down minerals into residual frameworks, and ions in solution. Even the most resistant minerals can eventually be decomposed chemically, given suitable conditions over a long enough time. The products of weathering are 1 unaltered primary minerals and rock fragments; 2 new secondary minerals; 3 ions in solution. For example, Table 2.1 shows the breakdown of a rock (granite) and a mineral (orthoclase). Mineral stability at the earth’s surface is the reverse of Bowen’s reaction series for the crystallization of igneous rocks (Fig. 2.1). Minerals formed at high temperatures and held together by mainly ionic bonds, can be ionized, hydrolyzed, etc., and chemically decomposed much more readily than minerals formed at lower temperatures and held together by mainly covalent bonds.

14 Chapter 2 Table 2.1 Chemical breakdown of granite and orthoclase. Primary constituents Minerals

Weathering products Colloids

Secondary minerals

Primary persisting

Granite Alkali feldspar Si, Al Clays Quartz Si Mica Si, Al Clays Ferromagnesium minerals Si, Al Clays, Fe oxides Orthoclase KAlSi3O8 + 2H+ + H2O = 2K+ + Al2Si2O5(OH)4 + SiO2 (orthoclase) (kaolinite) (colloidal silica + quartz)

Soluble ions Na, K

Quartz some mica

Ca Fe, Mg

The formula of kaolinite could be written as Al2O3.2SiO2.2H2O to emphasize the presence of water.

Fig. 2.1 Bowen’s reaction series and temperature of mineral formation.

Chemical reactants vary greatly in their effects depending on the atmosphere, hydrosphere, and rock types. The main reactions involve water in some way: direct solution, hydrolysis, hydration, oxidation, and reduction. Direct solution of primary minerals without chemical change is rare, although silica phytoliths often simply dissolve in groundwater. Usually solution removes ions produced by other processes. In hydrolysis, water reacts with silicate minerals to produce usually clay minerals, ions, and some soluble

silica; for example, the breakdown of orthoclase shown in Table 2.1. Less complex, more soluble minerals produce ions only; for example, the breakdown of calcite: CaCO3 + H2O + CO2 = Ca 2+ + 2HCO3(calcite)

(carbonic acid)

(calcium cation)

(bicarbonate anion)

(2.1)

In carbonation, which is the opposite of hydrolysis, carbon dioxide ions are added to the minerals. The reversible reaction (2.1) goes to the left. Carbonation causes precipitation of carbonate and cementation of

Weathering 15 soil horizons, and often occurs during evaporation or photosynthesis. In hydration, minerals absorb water, expand, and fall apart. In oxidation, oxygen ions (usually dissolved in water) are added to the mineral (or hydrogen ions are removed). The oxidation process often transforms soluble minerals into insoluble minerals; for example, the oxidation of ferrous to ferric iron, which creates the brown and red colors in soils. In reduction, hydrogen ions are added to the minerals (or oxygen ions are removed). Ferric oxides may be reduced to ferrous oxides, changing soil colors to blue or green; and then be removed in solution, bleaching the soil horizon to a gray color. Oxidation and reduction of iron compounds are responsible for many of the colors of sedimentary rocks.

2.2 Rates of weathering Rates of weathering depend on a host of factors, such as rock type and mineral stability at the earth’s surface, climate and slope, and on chemical reactants present in the atmosphere and hydrosphere. The proportion of physical to chemical weathering and their rates vary due mainly to rock type, temperature and availability of water. Rock types control weathering by their physical and chemical reaction to climate. Igneous rocks tend to form resistant masses in cold and dry climates where physical breakdown is faster than chemical breakdown. Metamorphic rocks vary depending on their structure and mineral stability. Sedimentary rocks also vary. Massive limestones may form high jagged mountains in high latitudes where the physical breakdown of adjacent layered sandstones and shales is faster. However, similar limestones may form valleys in wetter and warmer lower latitudes where chemical solution is more important. Temperature controls weathering because the rate of chemical reactions doubles with every 10 °C rise. So, chemical weathering is faster in hot than in cold areas. Water is required for most chemical reactions, so chemical weathering is faster in wet than in dry areas. As a gross oversimpliﬁcation we can distinguish three extreme climatic environments on land that control the type and rates of weathering: warm and wet areas; warm and dry areas; and cool and mostly wet areas (Fig. 2.2).

Warm and wet areas, with high to moderate rainfall, have running water present at all times. These areas are dominated by chemical weathering, with thick, mature leached soils and lush vegetation. The high temperatures and throughﬂow of water promote rapid chemical reactions and removal of soluble products. In such areas, massive unstable rocks such as limestone form lowlands, and hills are low and rolling and covered in vegetation (Fig. 2.3a). Even rapidly rising mountains, such as those of the eastern Himalaya, may show rounded proﬁles covered in vegetation. Modern examples are lowland tropical areas such as the Amazon Basin and much of Southeast Asia. Warm arid and semi-arid areas, with low and sporadic rainfall, have running water present only intermittently. These areas are characterized by mixed physical and chemical weathering, with alternation between solution and re-precipitation during wet and dry periods. Most also have both diurnal (dew) and seasonal cycles of wetting and drying, which can lead to the formation of calcretes and other chemical precipitates in submature soils. The relatively high temperatures allow rapid chemical weathering at those places and on those occasions when water is present. Massive unstable rocks such as limestone form hills in such areas, since physical weathering can break down fractured rocks more rapidly (Fig. 2.3b). Modern examples are the deserts and semi-deserts of the American Southwest, northern Africa (the Sahara), and Central Asia. Cool glacial and periglacial environments, with variable rainfall, have running water only intermittently present due to the low temperatures. These areas are dominated by physical weathering, including freeze–thaw cycles. Even when water is ﬂowing, the low temperatures inhibit chemical reactions. Massive unstable rocks can form mountains in such areas, with fractured intervening rocks forming valleys (Fig. 2.3c). Modern examples are the mountains of the American Northwest, northern Eurasia, Tibet, and Antarctica. If weathered material is not immediately removed by one of the agents of transportation, then a weathered mantle or soil accumulates. It is rather unusual for particles to be weathered, transported, and ﬁnally deposited without at least spending some time in a soil proﬁle. So understanding soil processes and formation is important in interpreting sediments and sedimentary rocks.

16 Chapter 2

Fig. 2.2 Regions of differing climatic regimes (after Weyman & Weyman 1977, ﬁg. 2). (Reproduced with permission of HarperCollins.)

2.3 Soil formation Where slopes are not too steep, soils accumulate. The soil formed is controlled by the initial parent composition and the nature and rate of weathering. These, in turn, depend on the original rock type, temperature, availability of water (summarized by climate), and time. Soils are important in stratigraphy: they modify and change the materials produced by simple weathering; they produce new particles and modify groundwaters; they can provide information of sedimentary processes and environments; and they can be important stratigraphic horizons in their own right (Taylor & Eggleton 2001). Soil formation involves three processes: (i) the production of inorganic soil material by the weathering of the parent bedrock or sediment; (ii) the incorporation

of organic matter formed by the decomposition of plant and animal tissues; and (iii) the reorganization of these components by aggregation and translocation to form SOIL HORIZONS. These processes produce a bewildering variety of soils, which nevertheless usually show (at least roughly) the well-known threefold division into A horizons of organic accumulation and leaching; B horizons of accumulation; and C horizons of partially altered bedrock (Fig. 2.4). Weathering is often the only one of these three processes that is covered in stratigraphy texts. However, equally important are the processes that transform organic matter and cause reorganization within soils. These processes not only produce new minerals and aggregates but can even produce new rocks and particles that will undergo further weathering and transport. The

Weathering 17 (c)

(a)

(b)

Fig. 2.3 Examples of the main regimes: (a) warm and wet; (b) warm and dry; and (c) cool and wet.

18 Chapter 2 (a)

(b)

Fig. 2.4 Highly weathered tropical ferrasol: (a) an example of a highly chemically weathered tropical ferrasol; (b) typical proﬁle of a tropical ferrasol (from Fitzpatrick 1980, ﬁg. 1.2). (Reproduced with permission of Pearson Education Limited.)

rocks and particles produced by soil processes may easily be confused with those produced in other ways. Many plants produce sand-sized colloidal silica particles called phytoliths (which are indistinguishable from the small chert particles produced during diagenesis of limestones). The massive silica beds (silcretes) that can precipitate in stable semi-arid soils are difﬁcult to distinguish from cherts formed by the replacement of limestones (Fig. 2.5). The nodular and massive carbonates (calcretes) that form in semi-arid soils are difﬁcult to distinguish from marine limestones. Thus, some supposed Cretaceous marine limestones of Central India are now reinterpreted as thick soil calcrete horizons, requiring major changes in Indian Cretaceous stratigraphy and paleogeography (Brookﬁeld & Sahni 1987). Furthermore, silcretes and calcretes may then be weathered and transported; and such particles may be difﬁcult to identify.

Unless you can recognize that these particles were produced in soils, you may infer a non-existent limestone and/or chert source. Organic matter is decomposed by soil micro-organisms, dominantly bacteria and fungi, but sometimes in the guts of larger organisms such as earthworms, producing humus. Mixing of this partially decomposed organic matter with altered rock and minerals forms soil A horizons. Decomposition is favored by alkaline and neutral conditions, and both larger organic and bacterial activity drop sharply with increasing acidity. In neutral or alkaline conditions, such as under grassland, the complete breakdown of organic matter is accompanied by reworking of the upper soil layers, producing a mull humus of almost completely decomposed organic material in a well-aerated spongy fabric of clay minerals with silica phytoliths, held together with clay minerals, sesquioxides, and polysaccharides. With increasingly

Weathering 19 (a)

(b)

Fig. 2.5 Silcrete, Lake Eyre, Australia: (a) massive silcrete; (b) reworked silcrete.

20 Chapter 2

Fig. 2.6 Soil catena in a tropical river valley and soil proﬁles (from Knapp 1979, ﬁg. 3.10). (Reproduced with permission of HarperCollins.)

acidic conditions, such as under a beech wood, bacteria and larger animals become scarcer, and organic material is only partially broken down, producing a moder humus of partially decomposed plant material mixed with mineral grain aggregates and arthropod droppings. A thin plant litter remains at the surface. With truly acidic conditions, such as under coniferous forest or heath, the quantities of micro-organisms decrease further and larger animals become very rare, producing a mor humus of only slightly decomposed plant material which may accumulate to a considerable thickness and show a clearly deﬁned layer structure. With extreme waterlogging and maximum acidity, microorganism activity is inhibited to such an extent that plant tissues are only slightly decomposed, producing peat. Reorganization involves two contrasting types of process: those tending to aggregate materials into stable masses, and those tending to move or translocate material to different places. Both tend to emphasize the horizon boundaries of soils. Aggregation occurs when soil

passing through the intestines of earthworms and other soil animals gets cemented by polysaccharides; aluminum and iron oxides precipitate as coatings on, and bridges between, soil particles in wet climates; and dissolved bicarbonates precipitate in the same way in semiarid and arid climates. Translocation separates soils into distinct layers, which may be emphasized by local aggregation. In soils with little organic matter and insufﬁcient cations to stabilize ﬁne particles, material gets removed from an eluvial (E) horizon and washed down into an illuvial (B) horizon to accumulate and precipitate as grain coatings and bridges. These can eventually develop into impermeable levels, as in calcrete. The effect of all these processes is to determine the nature of the material supplied for transportation. For example, in cold climates with rapid erosion, granite will supply rock fragments and unstable minerals with little clay or soluble ions. In dry, hot climates with slow erosion, granite will supply a mixture of rock fragments and minerals, though at least some of these will be broken down chemically to clays and ions in solution.

Weathering 21

Fig. 2.7 Subaqueous soil formation: (a) depth of burrowing related to oxygenation of sediments (after Wignall 1993, ﬁg. 1 (Reproduced with permission of the Geological Society.); (b, c) variation in intensity and depth of burrowing in one benthic community (after Bromley 1990, ﬁg. 6.10). (Reproduced with permission of Harper Collins.)

In wet, hot climates with slow erosion and deep soil development, granite will supply only quartz, clay minerals, and ions in solution. Lastly, local conditions cause great variations in the nature and development of soils even on the same bedrock. A sequence of topographically related soils on the same bedrock is called a SOIL CATENA. In the soil catena shown in Fig. 2.6, from a seasonally wet tropical climate, the valley slopes are free-draining, thoroughly leached during the rainy season, and oxidized in the dry season: soils tend to be coarse textured and ferruginous. Downslope, under savannah vegetation, less leaching takes place: soils tend to accumulate sesquioxides and clay minerals forming laterites. At the base of slope, under lush vegetation, waterlogging forms organic-rich, gleyed soils.

2.4 Weathering and soil formation under water Except that the surfaces are constantly bathed in water, underwater weathering and soil formation are the same as on land. Weathering depends on rock type, temperature, water chemistry, and the biota. Basalt is altered by hydrothermal circulating systems at mid-oceanic ridges. Soft sediments get churned up by burrowing organisms, and become aggregated into fecal pellets as in subaerial soils (Frey & Wheatcroft 1989). These processes are controlled by the physical and chemical properties of the sediment and overlying water, and by the type of organism capable of living in these conditions (Fig. 2.7). Underwater weathering and soil formation are as important as on land in controlling sediment characteristics.

3 Sediments and sedimentary rocks

3.1 Transportation and deposition 3.2 Clastic sediments and sedimentary rocks 3.3 Chemical and biochemical sediments and sedimentary rocks For the main sediment and sedimentary rock classes, we need to know how and by what agents they get transported and deposited; how they change both during and after deposition; and what classiﬁcation will reﬂect, simply and effectively, how they formed. This chapter is a very brief reminder of the essentials of sediment and sedimentary rock formation and classiﬁcation. You should refer to sedimentology texts for further information; for example, Allen (1985), Blatt et al. (1972), Blatt (1982), Leeder (1999), McLane (1995), Prothero and Schwab (1996), Selley (1985), Tucker (1982, 1991), and Tucker and Wright (1990).

3.1 Transportation and deposition During transportation, sedimentary particles and solutions are modiﬁed by a variety of physical and chemical processes which eventually produce clastic, biochemical, or chemical sediment. Transportation is driven by gravity. Gravity by itself pulls solid material downhill to form scree and talus slopes and rock glaciers, though even here water often

lubricates the contacts, reducing friction and causing movement. Gravity also acts on heavy water-saturated soils to pull them downhill as mass ﬂows (Fig. 3.1). It acts on water in streams, pulling it and its dissolved ions and enclosed particles downhill. Lastly, gravity controls the movement of ice in glaciers and acts on displaced air and water masses to cause winds, waves, currents, and tides. Solid materials such as weathered rocks, minerals, and newly formed soil minerals and biochemical constituents are transported and deposited as particles to form clastic sediment (or bioclastic sediment), composed of gravels, sands, silts, and clays. Here, transportation depends on the physics (e.g. the viscosity and velocity) of the transporting ﬂuid (if any). Ions in solution are passively transported by groundwater and streams to be precipitated by organisms to form biochemical sediments, or by changes in concentration to form chemical sediments. Here, transportation depends on the concentration of ions and the ﬂuid chemistry. Deposition is driven by physical or chemical changes in the depositing medium. Clastic sediments are deposited when grains slow down, or when the capacity or competence of a moving mixture of grains and

Sediments and sedimentary rocks 23

Fig. 3.1 Mudﬂow on Mt Rainer (courtesy USGS/Cascades volcano observatory).

Fig. 3.2 Hjulstrom’s diagram showing the critical velocity for movement of quartz grains on a plane bed at 1 m water depth. The shaded area is experimental scatter (after Sundborg 1956).

ﬂuid decreases. Deposition of biochemical sediments initially occurs when organisms extract chemicals from solution or the atmosphere, though the particles formed from these organisms are then worked on in the same way as clastic particles. Deposition of chemical sediments occurs when solute concentrations reach saturation point, either by evaporation or by changes in Eh or pH of the water. Obviously all three processes can occur together. Of course, most clastic and many biochemical sediments are continually being deposited, eroded, and redeposited during the journey to their ﬁnal resting place. Silts are the easiest particles to erode and transport, since clays stick together electrostatically once deposited (Fig. 3.2). The amount of information about transportation and deposition that is preserved varies among clastic, biochemical, and chemical deposits. CLASTIC SEDIMENTS move from weathering proﬁle to depositional site as solid particles. The particles show how the parent rock has been modiﬁed during weathering and transport, how the sediment was deposited, and the processes that turned the sediment into rock. Composition, grain size, sorting, roundness, depositional texture, and diagenetic changes reﬂect these factors. BIOCHEMICAL sEDIMENTS are precipitated from solutions by chemical changes caused by organisms. The deposits show what organisms formed the sediment, the way in which these sediments have been modiﬁed by transport, how the sediment was deposited, and the processes that turned the sediment into rock. Taxonomic composition, grain size, grain shape, sorting, roundness, depositional texture, and diagenetic changes reﬂect these factors. However, unlike clastics, biochemical rocks do not reﬂect their weathering source or solute transport history.

24 Chapter 3

Fig. 3.3 Maturity concept for clastic sediments (from Blatt 1982, ﬁg. 4.8).

Furthermore, recrystallization and replacement often causes problems due to alteration of original features such as depositional texture. For example, carbonate particles can be aggregated to form ooliths, grapestones, and fecal pellets, and broken to form rock fragments. These particles then behave as detrital grains. CHEMICAL SEDIMENTS show simply the geochemical conditions of precipitation from water, including lakes, seas, groundwater, and porewaters. Composition may be the only signiﬁcant feature preserved because of the ease of recrystallization.

3.2 Clastic sediments and sedimentary rocks Physicists and engineers have long studied the behavior of solid grains both en masse and in various ﬂuids (see Allen 1985). Gases and liquids are both ﬂuids since they lack shear strength. Thus the behavior of grains in liquids and gases is similar, accounting for the difﬁculty in distinguishing some wind-deposited sedimentary rocks from water-deposited ones (Selley 1982). Most clastic sediments are transported and deposited in water. After in-situ weathering and during water transport, particles are further broken down physically by collision with each other and the underlying bed. They are bathed in water, continue to break down chemically, and get progressively smaller, better rounded, and better sorted. The result is that the original soil mixture

separates into distinct sizes and further decomposes during transport: it changes from a mixture of unstable, angular, poorly sorted fragments into a series of stable, rounded, well-sorted sedimentary particles. This (ideal) change is expressed in the maturity concept (Fig. 3.3). IMMATURE SEDIMENTS are those least modiﬁed from the parent material. They have variable grain sizes, and consist of a wide variety of poorly rounded mineral and/or rock fragments, including those easily broken down physically and chemically. MATURE SEDIMENTS are those most modiﬁed from the parent material. They have a restricted range of grain sizes, and consist of moderately to well-rounded mineral or rock fragments of limited variety, often only quartz and/or clay minerals. All variations exist between these extremes. In water ﬂows, at least, it is useful to distinguish extraclasts from intraclasts. Extraclasts are derived from outside the depositional basin, while intraclasts are derived from inside the basin, often from soil nodules. Deposition of clastics takes place either by freezing of mass ﬂows, by the gradual build-up of traction carpets by migrating bedforms, or by settling out of suspension. None of these are mutually exclusive. Underwater mass ﬂows often mix with the surrounding water, changing from viscous to ﬂuid ﬂows, with the resulting depositional mechanism changing from freezing to traction. As water ﬂows slow down, depositional mechanisms change from traction to suspension; for example,

Sediments and sedimentary rocks 25

Fig. 3.4 Characteristics of simple and complex craters (from Mark 1987, ﬁg. 12.6).

climbing sand ripples are often draped by mud laminae. How much a clastic sediment is modiﬁed during transportation and deposition depends on the agent; whether this is an impact or an explosion (ejecta blanket), a mass ﬂow (scree, debris ﬂow), a water ﬂow (stream, lake, or ocean), an ice ﬂow (glacier), or an air ﬂow (wind). 3.2.1 Impact and volcanic processes Extraterrestrial impacts and volcanoes hurl material outwards from a central point. This material both ﬂows outwards and settles out of suspension, forming chaotic boulder beds, gravels, and graded sands and silts. It can then be (and often is) further moved by the usual agents of gravity, water, ice, and wind. IMPACT CRATERS and their deposits were recognized by very few people before the 1970s (the most famous being the Sudbury Basin by Dietz in 1963) (Deutsch et al. 1995). Not until the Apollo missions proved that the moon’s craters were caused by impacts rather than volcanism, did a serious search for impacts on earth begin. Over 120 impact craters are now recognized on the earth, ranging in age from several thousand to two billion years old, although most are younger than two hundred million years, since craters get destroyed over time by erosion or sea-ﬂoor spreading, or obscured by sediment (Grieve 1990). Major impacts are plausibly used to explain great events in earth history, such as the mass extinction at the end of the Cretaceous (Alvarez et al. 1980).

There are two basic forms of impact crater (Fig. 3.4). Simple craters are 2–4 km across and have a simple heavily fractured and brecciated bowl, surrounded by ejecta and partially ﬁlled with breccia and impact meltrock lenses. Complex craters can be one hundred times wider than they are deep. A central uplift is surrounded by rings marking the collapse of transient earth waves; in between the uplift and the rings are melted, brecciated, and shocked materials (Spudis 1993). Impact craters can be confused with other geologic features due to explosions and collapse, especially volcanic calderas (small) and cratonic basins (large). However, impacters hit the earth at an average speed of about 25 km/s, and vast amounts of energy are released in a small area in a very short time. No internal planetary process can produce the shock pressures of some impacts: pressures are far too low in even the most stupendous volcanic explosions. Impact pressures exceed 100 GPa (106 ¥ atmospheric pressure), and temperatures can reach several thousand degrees centigrade. Diagnostic features of impacts include shatter cones, multiple shock lamellae in mineral grains, high-pressure minerals such as diamond (from carbon) and stichovite (from quartz), diaplectic glasses, and impact melts (for details, see French 1990). Impacts generate rapidly moving debris avalanches and ballistic fallout. Debris strikes the ground causing ground-hugging debris avalanches to hurtle outwards as low-density surges (Melosh 1989; Rampino 1994). Such deposits may be confused with volcanic or glacial

26 Chapter 3

(c)

Fig. 3.5 (a) Cross-section through avalanche deposits of the Popigai crater (after Masaitis 1994, ﬁg. 7) (Copyright © 1994. Reproduced with permission of the publisher, the Geological Society of America, Boulder, Colorado, USA.); (b) diagram and (c) photo of breccia with impact melt fragments, Ries impact, Germany.

diamictites, debris avalanches, and debris ﬂows (Avermann & Brockmeyer 1992). For example, the coarse breccias from the 80 km diameter, 220 million year old, Puchezh-Katunki crater were originally used as evidence for Mesozoic glaciation (Oberbeck et al. 1993). However, details of the deposit, such as its general disorganization, impact melt glass fragments, and high-pressure minerals are distinctive (Fig. 3.5). Impacts into shallow seas should form well preserved and easily recognizable impact facies – keep a look out for them! Volcanoes form craters and radially ﬁning ejecta with interlayered rocks crystallized from melts, just like extraterrestrial impacts. Unlike impacts, volcanoes are concentrated in lines or groups, related to plate bound-

aries or mantle plumes, and their deposits are well studied (Cas & Wright 1987). Gentle basic eruptions form ﬁssure and shield volcanoes dominated by easily recognized massive or pillowed basaltic lavas. More explosive intermediate and acidic lavas form cones dominated by fragmental deposits interbedded with rarer lavas. These volcanoes may also blow up spectacularly, forming collapse calderas surrounded by debris from the explosion (Fig. 3.6). Since large volcanic explosions hurl lava, pyroclastics, and country rock into the air, the resulting deposits may be difﬁcult to distinguish from impact ejecta (Fig. 3.7). However, volcanic ejecta will not have multiple shock lamellae, high-pressure minerals, diaplectic glasses, or impact melts. In addition, volcanic eruptions and

Sediments and sedimentary rocks 27 explosions are repetitive: old volcanoes become inactive and new ones develop in the same belt. The result is a succession of diverse and interbedded volcanic and volcaniclastic facies (including explosion breccias) modiﬁed by normal sedimentary processes. 3.2.2 Mass ﬂows to water ﬂows Loose material on slopes can creep slowly downhill as scree or can move more rapidly as mass ﬂows or in water ﬂows. It is convenient to consider mass ﬂows and water ﬂows together as they form a continuum from debris avalanches through debris ﬂows and hyperconcentrated ﬂows to dilute water ﬂows (Fig. 3.8). Ice and air are more distinct. DEBRIS AVALANCHES are inertial, turbulent granular ﬂows caused by large landslides in which the particles

Fig. 3.6 Flooded caldera, southern part of St Kitts island, Caribbean.

are simply lubricated by the ambient ﬂuid (air on land, water underwater), which readily escapes from between the grains as energy is lost, and in which the larger clasts can impact readily during movement. They may move very slowly, as in rock glaciers, or very rapidly, as in some catastrophic slope failures, and may be enormous (Bugge et al. 1988). Deposits are characteristically nonerosive, massive, coarse, very poorly sorted, heterogeneous framework breccias in which individual clasts may be highly fractured and with an unsorted inﬁltrated matrix (Fig. 3.9). DEBRIS FLOWS are mixtures of sediment and ﬂuid, in which the larger clasts are supported by matrix strength and ﬂuid buoyancy as well as by the grain dispersion forces of debris avalanches, and in which the ﬂow is generally viscous. A distinctive characteristic is that the supporting ﬂuid does not mix with the ambient ﬂuid. Thus, debris ﬂows on land stay cohesive and remain concentrated from mobilization to depositional site: they may rapidly “freeze” and dewater as they lose energy downslope. Underwater debris ﬂows may slowly absorb water, becoming more ﬂuid downslope: they may deposit coarser material as competence decreases and change into hyperconcentrated ﬂows. Both subaerial and underwater debris ﬂows deposit non-erosive, massive, poorly sorted, matrix-supported gravels, with possible internal lamination due to shearing, and alignment of clasts parallel to ﬂow (Fig. 3.10). HYPERCONCENTRATED FLOWS are intermediate between viscous, non-turbulent debris ﬂows, and dilute, fully turbulent water ﬂows. Smith (1986) deﬁned hyperconcentrated ﬂows as high-discharge ﬂows in which turbulence is not the only sediment-support mechanism and in which deposition does not occur en masse.

(a)

(b)

Fig. 3.7 (a) Volcanic debris ﬂows and (b) ash deposits, St Kitts, Caribbean.

28 Chapter 3

Fig. 3.8 Sediment support mechanisms and typical deposits of mass ﬂows to water ﬂows (from Smith & Lowe 1991, ﬁg. 1). (Reproduced with permission.)

Hyperconcentrated ﬂows are ﬂuid enough and turbulent enough to let differential settling form graded beds. At the viscous end, ﬂuid buoyancy and grain dispersive forces are dominant. “Fluid debris ﬂows” deposit graded pebbly mudstones (Fig. 3.11a) while granular sand “grain ﬂows” deposit massive or graded beds with water-escape dish structures (Fig. 3.11b). At the ﬂuid end, turbulence dominates and migrating bedforms deposit graded cross-bedded sands and silts. Dilute hyperconcentrated ﬂows, known as TURBIDITY CURRENTS, often show a characteristic vertical sequence of structures (known as BOUMA SEQUENCES after their discoverer) due to progressive deposition during gradual ﬂow deceleration (Fig. 3.12). In addition, turbulence in dilute ﬂows causes erosion of the substrate before deposition. Turbidity currents show a variety of structures produced by either particles or vortices impacting and eroding the underlying material (Fig. 3.13). Though subaqueous turbidity currents were ﬁrst used to explain graded deep sea sands, graded beds are

also produced from any decelerating turbid ﬂow, such as those produced by storms on shelves and by ﬂoods exiting river mouths (Nelson 1982). Dilute water ﬂow is fully turbulent except in a thin laminar layer near the bed. Turbulence is the main sediment-support mechanism and deposition occurs in a “grain by grain” fashion either from traction carpets moving in migrating bedforms or from suspensions settling directly out on the surface. Traction deposits have a minimum coarse silt to ﬁne sand size, since silt and clay once deposited from suspension remain within the laminar sublayer unless affected by bursting or turbulent eddies. Sand moves as sheets or in migrating bedforms depending on grain size and ﬂow power. Migrating BEDFORMS form a variety of cross-beds depending on sediment supply, rate of bedform climb and current–wave ﬂow. Sediment-saturated decelerating ﬂows deposit sediment in rapidly climbing bedforms passing into suspension deposits (Fig. 3.14). Unsaturated, sediment-starved or steady ﬂows deposit trains of

Sediments and sedimentary rocks 29

Fig. 3.9 Debris avalanche and scar, Baralacha La, Lahul, Indian Himalaya.

erosive bedforms (Fig. 3.15a). Fast ﬂows form parallel laminae with lineations caused by boundary-layer bursting (Fig. 3.15b). Depending on the depth of ﬂows, several bedform hierarchies can co-exist, due to several scales of ﬂow vector ﬁelds and boundary layers (Allen 1982). Water ﬂows in standing bodies of water such as lakes and seas often have oscillatory or wave components. Normally these form simple wave ripples, interference ripples, and combined wave-current ripples (Fig. 3.16a,b). Less common but larger storm, hurricane, and tsunami waves have complex surface patterns with bottom wave surges and current drift varying in direction and affecting deeper areas. Waning waves of these types may form large irregularly superimposed crossbeds called HUMMOCKY CROSS-STRATIFICATION (Fig. 3.16c). Regular intermittent and reversing ﬂows occur in many tidal environments. Traction deposition in sand bedforms alternates with silt and clay suspension deposition, forming thin ripple trains separated by clays. Depending on the traction/suspension ratio, various types of FLASER BEDDING may result (Fig. 3.17). These, of course, can consist of current, wave, or combined current–wave bedforms at various scales. Plotting the orientation of both erosional and depositional features gives the spread of ﬂow direction(s) and possible ancient current systems (Fig. 3.18), which, in turn, may help you to recognize ancient environments (Table 3.1). Mass ﬂows and water ﬂows account for most clastic deposits, but ice and air ﬂows are important in cold and/or arid environments, and locally elsewhere.

Table 3.1 Various water environments and dispersal patterns (from Tucker 1991, table 2.5). (Reproduced with permission of Blackwell Publishing Ltd.) Environment

Dispersal pattern

Fluvial

Paleocurrents reﬂect paleoslope and indicate provenance direction; unimodal pattern with small dispersion if low-sinuosity rivers, unimodal with larger scatter if high-sinuosity river or alluvial fan

Deltaic

Typically unimodal pattern directed offshore although marine processes (tidal and storm currents and waves) can complicate paleocurrent pattern

Shallow-marine shelf

Pattern can be complex and difﬁcult to interpret; bimodal pattern through tidal current reversals although tidal currents may be parallel or normal to shoreline; can be unimodal if one tidal current dominates; polymodal and random patterns also occur; complicated by wave and storm effects

Turbidite basin

Unimodal pattern common from turbidites, although may be downslope or along basin axis, or radial if on submarine fan. Contourites give paleocurrent pattern parallel to the strike of the slope

30 Chapter 3 (b)

(c)

Fig. 3.10 (a) Debris ﬂow characteristics (from Allen 1985, ﬁg. 9.24). (Reproduced with permission of HarperCollins.) (b, c) Triassic debris ﬂow, British Columbia.

3.2.3 Ice ﬂow (glaciers) Glaciers ﬂow plastically, carrying unsorted sediment on and within them (Fig. 3.19). Deposits directly from glaciers (TILLS) resemble those of debris ﬂows and impact and explosion ejectas. Unsorted material is smeared along the base of the ice, dumped at the terminus, or dropped from ﬂoating ice as massive, poorly sorted, pebbly mudstones (DIAMICTITES) (Fig. 3.20). Enormous shear pressures at the base of the ice may produce highly compacted sheared mudstones with strongly aligned pebbles, while glaciers melting in quiet water may deposit material showing crude stratiﬁcation and grading. Icebergs can ground, carving huge grooves and depressions into the lake or sea ﬂoor, as well as melt far from the glacier, dumping unsorted lenses in any other sediment. Glaciers grind and shatter quartz grains and cold conditions limit chemical weathering,

so glacial diamictite matrices often contain very immature silt and very angular quartz silt. These contrast with the less shattered, chemically rounded, more mature silts of debris ﬂows in arid regions. It may be very difﬁcult to tell glacial diamictites from debris ﬂows in cold wet areas (Fig. 3.21).

3.2.4 Air ﬂow (winds) Wind deposits resemble deposits from dilute water ﬂows, except that the lower viscosity and density of air limit transport to granule-size particles and below, and allow destructive impacts at lower grain sizes. Fine sand gets rounded and minerals with cleavages get broken faster. Threshold stresses vary more rapidly for given grain diameters in air than in water thus producing better sorting of individual laminae (Fig. 3.22). Wind ripples

Sediments and sedimentary rocks 31 (a)

(b)

(c)

Fig. 3.11 (a) Fluid debris ﬂow: graded conglomerate, Jurassic, British Columbia. (b) Detail of (a). (c) Grain ﬂow: coarse sandstone, Precambrian Jura Quartzite, Scotland.

Fig. 3.12 Bouma sequence and variation down-current (from Allen 1985, ﬁgs 12.24, 12.25). (Reproduced with permission of HarperCollins.)

32 Chapter 3 (a)

(b)

Fig. 3.13 Casts at the base of Ordovician sandstones: (a) vortex scours; (b) tool prod marks.

concentrate coarse sand at their crests (in contrast to water ripples). The very low angle of climb of wind ripples also leads to distinctive translatent strata (Fig. 3.23). The atmospheric boundary layer is also much thicker (over 1 km) than any equivalent water ﬂows (except those deeper than the outer shelf) allowing larger bedforms and more hierarchies to co-exist (Fig. 3.24). 3.2.5 Post-depositional modiﬁcations

Fig. 3.14 Supercritical climbing ripples in ﬁne sand, with a silt/clay drape (Quaternary, Ontario).

The diagnostic criteria that are used to infer how an ancient sedimentary rock layer was deposited may be modiﬁed by post-depositional processes when clastic sediments are turned into rocks. Post-depositional modiﬁcations start as soon as sediment is deposited. Usually, near-surface physical changes, such as loading, slumping, and modiﬁcation by organisms, are separated

Sediments and sedimentary rocks 33

Fig. 3.15 (a) Subcritical climbing ripples or dunes; (b) plane bed lamination in coarse sand (both from Harms et al. 1975, ﬁgs 3.1, 3.2). (Reproduced with permission of SEPM.)

from diagenesis which typically involves deeper physical and chemical changes, but of course these overlap. For example, clays mostly compact and lose water very near the surface. 3.2.5.1 Near-surface physical changes Loading occurs when denser sediments differentially sink into less dense sediments, forming structures such as convolute lamination, convolute bedding, and balland-pillow structures (Fig. 3.25). Frequently the denser layers are sands overlying less dense water-saturated clays that become ﬂuidized by earthquakes and other disturbances. Loading is often accompanied by a loss of porewater, which may lead to the formation of internal and surface structures such as sandstone dikes and sand volcanoes (Fig. 3.26). Slumping occurs when sediment layers move sideways as well as vertically. Often this occurs when ﬁne-

grained clastic sediments are deposited very rapidly on slopes, in places such as delta fronts and shelf edges. Slumps and slides may form folds and faults at any scale and cause total disruption or even homogenization of layered sediments (Fig. 3.27). It can be very difﬁcult, especially in orogenic belts, to decide whether some folds, faults, and melanges are due to sedimentary or tectonic processes. Various other minor features, such as ice wedges, rainprints, mudcracks, synaeresis cracks, and mineral pseudomorphs, may help in determining process. For example, mudcracks show at least occasional exposure to air, while halite pseudomorphs indicate evaporitic conditions. Modiﬁcations by organisms (BIOTURBATION) range from slight surface scratches to almost complete homogenization by burrowing organisms (Bromley 1990). Although such trace fossils (tracks, trails, burrows, and borings) destroy sedimentary fabrics and structures,

34 Chapter 3

(a)

Fig. 3.16 (a) Wave ripples, Recent tidal ﬂat, Scotland. (b) Block diagram of simple wave ripples (from Reineck & Singh 1973, ﬁg. 25). (Copyright (1973). Reproduced with permission of Springer-Verlag.) (c) Hummocky cross-stratiﬁcation (from Harms et al. 1975, ﬁg. 5.5). (Reproduced with permission of SEPM.)

Sediments and sedimentary rocks 35

Fig. 3.17 Classiﬁcation and block diagram of ﬂaser and lenticular bedding (from Reineck & Singh 1973, ﬁgs 164, 167). (Copyright (1973). Reproduced with permission of Springer-Verlag.) (b)

Fig. 3.18 (a) Various plots of current direction types (from Tucker 1991, table 2.5). (Reproduced with permission of Blackwell Publishing Ltd.) (b) Polymodal distribution of graptolites, Ordovician, Ontario, Canada.

36 Chapter 3

Fig. 3.19 Terminus of the Portage Glacier, Alaska.

they do give other information about the sediment and record the behaviour of organisms (Pemberton et al. 1992). Different suites of trace fossils have been used to characterize marine environments (Fig. 3.28). While trace-making organisms evolved through time, basic behaviors persist, and different trace-fossil suites tend to dominate different environments. Deposit feeders, which eat organic material in sediments, dominate the low-energy Cruziana ichnofacies; while suspension feeders, which ﬁlter material out of suspension, dominate the high-energy Zoophycos ichnofacies. 3.2.5.2 Diagenesis Diagenesis is all the physical, chemical, and biological changes that change a sediment after it has been deposited. Sedimentary characteristics, such as grain packing, porosity, and permeability, may be greatly altered by both physical processes such as compaction, and chemical processes such as cementation and alteration. Compaction occurs when grains are rearranged into denser conﬁgurations and ﬂuids are squeezed out of them. Modern clays may be deposited in very open water-supported frameworks with very high porosities of 80%. These porosities rapidly decrease to 40% as water is expelled during burial (Fig. 3.29a). Modern sands may be deposited in loose open frame-works with more than 50% porosity. These porosities decrease as grains are rearranged by pressure, sometimes bent and broken, and dissolved at pressure points (Fig. 3.29b). Cementation occurs at any time after burial and

during compaction. Common cements, such as silica and carbonate, may simply precipitate from groundwaters. Diagenetic histories of coarser clastics can sometimes be inferred from several generations of cements. Early nodular cementation sometimes fossilizes high porosities (and fossils in three dimensions). Calculated porosities for such nodules can approach 80%, showing that the carbonate was precipitated while the clays were essentially uncompacted (Fig. 3.30). Alteration is most serious when it modiﬁes the grain composition and the grain/matrix proportions. Many sandstones show authigenic growth of feldspars, and the grains of immature marine sandstones can react with sodium-rich porewaters, changing mixed calciumsodium plagioclase into almost pure sodium albite. If neither change is detected, weathering and source characteristics may be misinterpreted. Unstable mineral grains may break down to form a secondary clay matrix. Since most modern turbidite sands show graincontact textures, the ﬂoating textures of many ancient greywackes may be secondary (due to the breakdown of detrital volcanic rock and ferromagnesian mineral grains), and have nothing to do with depositional process. 3.2.6 Classiﬁcation The aim of any sediment classiﬁcation is to express, in a concise way, the characteristics of a sediment or sedimentary rock that are useful in showing how it formed. These classiﬁcations should be as simple as is possible for the purpose intended. Although, as Blatt (1982) noted, “several dozen classiﬁcations of sandstones . . . have been accompanied by new terms whose number and interpretation are rivalled only by the manufacturers of dry cereals and household detergents,” most studies need only a few basic terms. For clastic sediments, the basic features of composition, grain size, shape, roundness, sorting, and depositional texture and structure (packing, porosity, and permeability) are affected by diagenetic changes (compaction, cementation, solution, etc.) during transformation into sedimentary rocks. Grain composition is controlled by the composition and structure of the source rocks, by how these get weathered, and by how the grains get eroded, transported, deposited, and eventually changed to rock. Grain size is controlled by the energy level of the transport and deposition. Because settling velocity is directly proportional to the square root of the grain diameter, grain size intervals on a logarithmic (phi) scale are

Sediments and sedimentary rocks 37

Fig. 3.20 Till types related to their position in the glacier (from Hambrey & Alean 1992, p. 98). (Reproduced with permission of Cambridge University Press.)

(a)

(b)

Fig. 3.21 (a) Flow tills at a glacier terminus; (b) Quaternary lodgement till, Annan, Scotland.

38 Chapter 3

Fig. 3.22 Threshold stresses for water and air compared (from Allen 1985, ﬁg. 4.3). (Reproduced with permission of HarperCollins.)

(b)

(a)

Fig. 3.23 Eolian lamination: (a) inversely graded; (b) climbing translatent.

usually used (Table 3.2). Thus mechanical size analysis of sands can be compared with grain settling size analysis of sands, silts, and clays. Grain population statistics useful in interpretation (e.g. median, mean, sorting, and skewness) can be calculated directly from the raw data. Grain roundness is controlled by grain size, hardness, and rate of abrasion, because roundness is a result of impacts with other grains during movement (Fig. 3.31). Quartz is hard, resistant to both chemical and mechanical rounding, and usually taken as a standard for roundness. Generally, quartz particles larger than 5–10 mm are nearly always rounded, while those less than 0.1 mm are nearly always angular. In between, sand-sized particles are rounded only in air where particle impacts are not cushioned, and in water with high average kinetic energies and constant abrasion, such as on beaches.

Grain shape is inherited from the original mineral, rock, or organism. Tabular pebbles come from layered sediments, and spherical pebbles from massive rocks such as granites. Elongate quartz sand grains come from metamorphic rocks, while more spherical grains come from igneous rocks. Bioclastic grains can be many shapes depending on the shape of the original organisms and how they (and their cementing agents) have been modiﬁed during transport. Sorting is a measure of the spread or standard deviation of the grain-size distribution (Fig. 3.32). It is controlled by the sizes available at the source, by the agents of transport and deposition, and by the amount of time these agents have been operating. Depositional texture is controlled by the type of transport and deposition. FLOATING TEXTURES, with the grains embedded in and separated by ﬁne sediment,

Sediments and sedimentary rocks 39

Fig. 3.24 Bedform hierarchies in a complex barchan dune–draa system (from Clemmensen & Abrahamsen 1983, ﬁg. 19). (Reproduced with permission of Blackwell Publishing Ltd.)

Fig. 3.25 Ball-and-pillow structures, Upper Ordovician, Ontario.

were deposited by mass ﬂows (Fig. 3.33a). FRAMEWORK with the grains in contact, were deposited by ﬂuids (air, water, dilute particle–ﬂuid mixtures) capable of escaping from between the grains (Fig. 3.33b). Both are frequently hardened by the precipitation of

TEXTURES,

chemical cement (which is difﬁcult to see in ﬂoating textures). Both sorting and texture seen under the microscope may be misleading since thin sections do not show the correct size of many grains, and many grains in frameworks are apparently not in contact (see Fig. 3.34). Grain packing during deposition varies with the mode of deposition. Thus, deposits of debris avalanches and grain ﬂows are more loosely packed than deposits from dilute ﬂows (Fig. 3.34a,b). Unless cemented early, loosely packed sands are often squashed during compaction (Fig. 3.34c). Silt- and clay-sized sediments are usually too ﬁne to study petrographically. Mineralogical and chemical analyses are more useful, and these can then be compared with mineralogical and chemical analyses of sandstones. 3.2.7 Terminology Grain size, composition, sorting, texture, and roundness are used to classify clastic rocks. Conglomerates have rounded pebbles which have

40 Chapter 3 (b)

(a)

Fig. 3.26 (a) Sandstone dike, Precambrian; (b) sand volcanoes, Recent.

Fig. 3.27 Deformed varved clays, Quaternary, Ontario.

been transported some distance; breccias have angular pebbles which have been transported only a short distance. Orthoconglomerates and orthobreccias have framework textures, and are transported by ﬂuids (usually water since air cannot normally move material coarser than sand). Exceptions are the poorly sorted volcanic orthobreccias (called agglomerates) thrown out by explosive volcanoes (and by impacts). Paraconglomerates and parabreccias have ﬂoating textures, and are transported by debris avalanches or mass ﬂows, which may be graded if the mass ﬂow has a low viscosity. Massive, poorly sorted paraconglomerates and parabreccias with silt–clay matrices are called diamictites. Sandstones (particularly coarse- to mediumgrained ones) are most useful for petrographic analysis since the grains are large enough to be identiﬁed in thin

Fig. 3.28 Ideal bathymetric distribution of some trace fossils (from Seilacher 1967, ﬁg. 3).

Fig. 3.29 Changes during compaction: (a) water loss in clays; (b) decrease of porosity in sands (from Tucker 1991, ﬁgs 3.6, 2.76). (Reproduced with permission of Blackwell Publishing Ltd.)

42 Chapter 3 section, but small enough for a representative sample to be collected in the ﬁeld. Like conglomerates, sandstones can have a cemented framework (arenites) or a ﬂoating texture (wackes). Four common types of sandstone are quartz arenites, lithic arenites (or litharenites), arkosic arenites, and greywackes (Fig. 3.35). Quartz arenites are mature, well-sorted frameworks of moderately to well-rounded quartz sand (Fig. 3.36a). They form only after extreme chemical weathering and/or long transport, and/or long reworking, or they are second-cycle sands, reworked from previously deposited mature sands or sandstones. Lithic arenites

Fig. 3.30 Early diagenetic carbonate concretions with clays compacted around them.

have more rock grains than feldspar. They tend to be poorly to moderately sorted, with subangular to moderately rounded grains (Fig. 3.36b). Most form in streams and deltas draining rising uplands. Arkosic arenites have more than 25% feldspar grains and are derived from granites and gneisses. They tend to be poorly sorted, with subangular to subrounded grains (Fig. 3.36c). Most form in streams close to the source. Greywackes tend to be poorly sorted with angular to subangular grains ﬂoating in a ﬁne-grained matrix. Felspathic greywackes have more feldspar grains than rock fragments; lithic greywackes have the reverse (Fig. 3.36d). Many greywackes were deposited by turbidity currents that mobilized immature sediments around magmatic arcs. The modern consensus is that the matrix of most greywackes is secondary and due to the decomposition of unstable grains (since modern turbidites do not contain much mud). However, this cannot be true of quartz wackes which are common off ancient passive margins. Some ﬂoating textured greywackes and sandy mudrocks (when not matrices to conglomerates) probably form either by wind blowing sand into mud or by organic mixing (bioturbation). Sandstone composition can help determine the tectonic setting (see Chapter 10). Mudstones and siltstones form almost two-thirds of the stratigraphic column, yet are poorly understood and inadequately studied. They are too ﬁne grained to study adequately or show much in hand specimen or thin section, and are easily altered after deposition

MUD

SAND

GRAVEL

Table 3.2 Standard grain-size scale for clastic sediments (from Blatt 1982, table 4.1). Name

Millimeters

Boulder Cobble Pebble Granule

4096 256 64 4

Very coarse sand Coarse sand Medium sand Fine sand Very ﬁne sand

2 1 0.5 0.25 0.125

Coarse silt Medium silt Fine silt Very ﬁne silt Clay

0.062 0.031 0.016 0.008 0.004 Ø

Micrometers

f -12 -8 -6 -2

500 250 125

-1 0 1 2 3

62 31 16 8 4 Ø

4 5 6 7 8 Ø

Sediments and sedimentary rocks 43

Fig. 3.31 Grain roundness of 0.5 mm quartz sands: (a) river sand from schist – angular grains; (b) beach sand from sandstones – dominantly moderately rounded grains; (c) desert sand – dominantly very well rounded grains (from Greensmith 1978, ﬁgs 6.2, 6.3, 6.5). (Reproduced with permission of HarperCollins.)

Fig. 3.32 Standard sorting images. Numbers on divisions are standard deviations dividing classes (from Blatt 1982, ﬁg. 4.3).

Fig. 3.33 (a) Floating texture, sandy mudstone; (b) framework texture, arkosic sandstone (from Greensmith 1978, ﬁgs 5.1, 5.6). (Reproduced with permission of HarperCollins.)

44 Chapter 3

Fig. 3.34 Packing and porosity: (a) loosely packed, highly porous (30%) beach sand; (b) moderately packed, moderately porous (15%) river channel sand; (c) compacted (by burial) arkosic sandstone with low porosity (<5%). Fields of view 3 mm (from Williams et al. 1954, ﬁgs 94, 95). (© 1954 by W.H. Freeman. Used with permission.)

Fig. 3.35 Petrographic classiﬁcation of sandstones (from Tucker 1991, ﬁg. 2.57). (Reproduced with permission of Blackwell Publishing Ltd.)

Sediments and sedimentary rocks 45

Fig. 3.36 Sandstone types in thin section: (a) quartz arenite; (b) lithic arenite; (c) arkosic arenite; (d) greywacke (lithic). Field of view is 2 mm (from Williams et al. 1954, ﬁgs 98, 100, 104, 108). (© 1954 by W.H. Freeman. Used with permission.)

(Fig. 3.37). They require chemical analysis and X-ray identiﬁcation of minerals for interpretation (Weaver 1989). Mudstones (and shales) consist of clay minerals (kaolinite, montmorillonite, illite, chlorite, or interleaved mixed-layer clays). The relative proportions of these clays minerals are strongly dependent on climate and weathering regimes, and so may be used to determine source characteristics (see Chapter 11). The

main problems are diagenetic alteration to illite, and increasing biological control on weathering over time. Siltstones consist of small mineral fragments, usually quartz, since the small size of the grains means that chemical breakdown is very rapid. Siltstones with unstable minerals (loess) were reworked from glacial deposits by katabatic winds during the Quaternary and form some of the most fertile (if unstable) soils in

46 Chapter 3

Fig. 3.37 Mudstones: (a) ﬁreclay with authigenic kaolinite; (b) shale with compressed clay minerals. Fields of view are 3 mm and 0.5 mm respectively (from Willliams et al. 1954, ﬁg. 111). (© 1954 by W. H. Freeman. Used with permission.)

Fig. 3.38 Classiﬁcation of clastic carbonates (from Greensmith 1978, ﬁg. 8.1). (Reproduced with permission of HarperCollins.)

northern mid-latitudes. Ancient loess must be common but has yet to be recognized. Clastic carbonates formed by the physical erosion of pre-existing limestones are classiﬁed by size in the same way as clastic rocks (Fig. 3.38). However, clastic carbonates have special problems. Often, you can neither work out the origins of the sand-sized carbonate grains, nor distinguish the intraclasts from the extraclasts. Furthermore, they often recrystallize, destroying the original fabric. Primary carbonates, formed by biochemical and chemical precipitation from ambient waters, are easier to interpret.

3.3 Chemical and biochemical sediments and sedimentary rocks Chemical sediments are precipitated when dissolved salts are concentrated by evaporation, or when chemical changes cause insolubility. Their formation is the reverse of chemical weathering. Particles are built up

rather than broken down. Only a few primary chemical precipitates, such as some carbonates, cherts and evaporites, are common as sediments. Of these, most carbonates and many cherts are actually biochemical sediments precipitated by organisms. Both chemical and biochemical sediments are widely distributed in soil, freshwater, marine, and hypersaline conditions, from poles to tropics and from shallow to deep water. 3.3.1 Carbonates Carbonates are the commonest chemical precipitates in both marine and continental waters. Even though most carbonates are precipitated either directly or indirectly by organisms, you still need to study the inorganic chemistry of carbonate precipitation. First, because organisms often catalyse or promote carbonate reactions, using the resulting precipitates to form supporting or protective structures. Mud-sized whitings of loose carbonate needles in the surface layers of the oceans may be due to photosynthetic changes in carbon dioxide

Sediments and sedimentary rocks 47 concentrations rather than to entirely inorganic chemical reactions (Shinn et al. 1989). Second, because unless you know the variety of natural carbonate deposits produced, you may misinterpret as organic what is really inorganic, and get a depositional process and environment wrong (Scholle 1978). For example, ooids (small snowball structures of calcium carbonate) often form in agitated marine waters by algal or bacterial precipitation, but regular ooids also form inorganically in caves. Unfortunately, there is a tendency to separate the study of continental (freshwater and groundwater) carbonates from marine (salt water) carbonates; some texts ignore freshwater and continental carbonates. Summaries of carbonate sediments and rocks can be found in Bathurst (1971), Flugel (1982), Milliman (1974), Reeckmann and Friedman (1982), Scofﬁn (1987), Tucker and Wright (1990), and Wilson (1975). 3.3.1.1 Mineralogy, isotopes, and chemistry It would be nice (but boring) if the system metal ion–H2O–CO2 were simple, containing only one mineral, and with solubility varying only with temperature and pressure and with simple chemical reactions. But this is not the case. There are three low-temperature polymorphs of calcium carbonate and one calcium– magnesium mineral whose inorganic mineralogy, chemistry, and isotope ratios are controlled by the ion concentrations, temperature, carbon dioxide, and oxygen levels of the ambient water. Furthermore, the mineralogy, chemistry, and isotope contents of organic carbonates are additionally inﬂuenced by physiology, taxonomy, and ecology. Mineralogy Calcium carbonate (CaCO3) has three lowtemperature polymorphs: CALCITE (trigonal), ARAGONITE (orthorhombic), and vaterite (hexagonal). Vaterite is metastable, of low density, and unimportant (though some ascidians precipitate it). So, calcite and aragonite are the only important primary carbonate precipitates. Experiments show that calcite is the only stable phase of CaCO3 at normal earth surface temperatures and pressures (Dickson 1990). The calcium–magnesium carbonate, DOLOMITE ([Ca,Mg]CO3), is an early replacement rather than a primary precipitate. Extrapolation of phase boundaries determined in the CaCO3–MgCO3 system above 400 °C shows that only magnesite, dolomite, and calcite with a few percent mol MgCO3 should be stable under earth surface conditions. Nevertheless, although marine oozes and some lake carbonates have the

predicted, thermodynamically stable calcite, many others have metastable aragonite. Calcite (the lower temperature, lower pressure form) is stable and occurs in two forms: high-magnesium calcite (11–20 mol% MgCO3) and low-magnesium calcite (<5 mol% MgCO3): intermediates are rare. The MgCO3 content of magnesian calcite precipitated from sea water is controlled by temperature and the CO2- concentration of the water, although the actual values depend on the taxonomy of the organisms (Mucci 1987). Magnesian calcites formed at low temperatures are metastable and are ultimately converted to a stable assemblage of low-magnesian calcite and dolomite (Mackenzie et al. 1983). The conversion often leaves the original texture unaltered at the microscopic level, so the conversion reaction is likely a microsolutionreprecipitation process. Fe2+ may substitute for Ca2+ and Mg2+ in small amounts (up to several thousand parts per million) during diagenesis to form ferroan carbonates. Aragonite (the higher temperature form) is stable at 25 °C only under pressures above 4 kbar. Despite this, many organisms secrete aragonite as their skeleton and aragonite is a common early cement. The Mg/Ca mol ratio in sea water (5 : 1) should favor precipitation of high-Mg calcite, but aragonite precipitates more readily, so Mg2+ must inhibit calcite precipitation (Morrison & Brand 1986). Sr2+ readily substitutes for Ca2+ in the aragonite lattice, the Sr/Ca ratio in aragonite being dependent on both the ratio in the solution and the temperature (Scofﬁn 1987). However, like the Mg/Ca ratio, the Sr/Ca ratio in aragonite is also taxonomically determined by physiology. The primary mineralogy of any carbonate sediment is determined not only by the physico-chemical conditions of the water, but also by the taxonomy of the carbonate-secreting organisms whose relative proportions have changed markedly through time (Table 3.3). Isotopes The isotopic composition of carbonates depends on the isotopic composition of the precipitating medium. In the CaCO3–CO2–H2O system, the stable isotopes of oxygen and carbon are 16O, 18O, 12C, and 13C, the ratios of which are controlled by environmental factors and the physiology of organisms. Isotopes are chemically precipitated in the same relative proportions they have in solution; however, they are often fractionated during biochemical precipitation, so that the isotope ratios systematically deviate from those of chemical precipitates. This deviation varies both among

48 Chapter 3 Table 3.3 Calcium carbonate mineralogy of some common organisms (from Tucker 1991, ﬁg. 4.7). (Reproduced with permission of Blackwell Publishing Ltd.) Organism Mollusca bivalves gastropods pteropods cephalopods Brachiopods Corals scleractinian rugose + tabulate Sponges Bryozoans Echinoderms Ostracods Foraminifera benthic pelagic Algae coccolithophoridae rhodophyta chlorophyta charophyta

Aragonite

Low-Mg calcite

x x x x

x

x

High-Mg calcite

Aragonite + calcite x x

(x) (x)

x x x

x x

x (x)

x x x x x

x

x x x

x x

x x

and within the main taxonomic groups of organisms. Stable carbon and oxygen isotopes of marine carbonates are reported as a deviation of the ratios 13C/12C and 18O/16O from arbitrary standards: SMOW (Standard Mean Ocean Water) for water; PDB (Peede Formation Belemnite) for carbonates. Deviations from equilibrium values are usually due to variations in temperature and salinity of the ambient sea water and/or biological fractionation. Formerly, it was often assumed that the overall isotopic composition of ocean water has remained essentially constant through time, and that the isotopic fractionation seen in ancient carbonates is due to temperature changes, salinity changes, and diagenetic alteration. However, paleotemperatures calculated from Paleozoic 18O/16O ratios of marine fossils are impossibly high when Recent marine salinities and sea-water isotope ratios are assumed to be the same – between 50 °C and 75 °C for the Devonian and Carboniferous from conodonts (Brand & Veizer 1983). Organisms frequently secrete carbonate with both carbon and oxygen isotope ratios in strong disequilibrium with ambient waters. Since most limestones are biogenic, isotope ratios in them are primarily affected by the taxonomy of the dominant organisms (Fig. 3.39). Carbon isotope ratios are mainly controlled by biochemical processes such as photosynthesis. Since the

Fig. 3.39 Carbon and oxygen isotope precipitation ﬁelds for some organisms (from Anderson & Arthur 1983, ﬁg. 1.23). (Reproduced with permission of SEPM.)

effect of temperature on carbon isotope ratios is very small, variations in marine carbon isotopes are thought to be worldwide and can be used as stratigraphic markers, even if the actual reasons for the variations are controversial (see Chapter 7).

Sediments and sedimentary rocks 49

Fig. 3.40 Major carbonate diagenetic environments in the subsurface (from Scofﬁn 1987, ﬁg. 8.1).

Chemistry The net result of the various equilibria equations for calcium carbonate in water is as follows: CO2 + H2O + CaCO3 = Ca2+ + 2HCO3Any processes that lead to a decrease in carbon dioxide and an increase in calcium ions in water drives this equation to the left, precipitating calcium carbonate; the converse drives the equation to the right, dissolving calcium carbonate. These processes include both physico-chemical processes (e.g. temperature and pressure changes, ion dilution, and concentration) and biological processes (e.g. photosynthesis and internal respiration). Increased temperature causes a loss of carbon dioxide and precipitation. Freshwater precipitation of carbonate occurs in large holes (caves), in small holes (soils), and in lakes. Fresh water dissolves calcareous rocks in places (forming karst and caves) and reprecipitates the carbonate

elsewhere (forming tufa, stalactites, nodules, and other cave and soil precipitates). The water table separates a vadose zone of intermittent saturation and drying from a phreatic zone of permanent saturation. Naturally this boundary ﬂuctuates, but it is usually a subdued duplicate of the topography above (Fig. 3.40). At the bottom of the vadose zone, just above the phreatic water table, the saturated water loses CO2 as it enters air-ﬁlled caves. This causes intense precipitation and the formation of characteristic cave precipitates of stalactites, stalagmites, dripstones, ﬂowstones, and cave pearls. In the phreatic zone, the holes and pores are ﬁlled with fresh water, which normally becomes more saline with depth, and eventually grades into stagnant connate (formation) water. Although freshwater carbonate precipitation is important in soils and during diagenesis, most carbonate particles form as biochemical precipitates in sea water supersaturated in CaCO2.

50 Chapter 3 Salt-water precipitation of carbonate occurs overwhelmingly in the sea (average salinity 36‰) and in saline lakes and soils. CaCO3 dissolves if CO2 increases due to pH changes or respiration, but it does not easily reprecipitate from salt water with a high Mg2+/Ca2+ ratio, so that sea water is often supersaturated. In sea water (with a molal Mg/Ca ratio of 5 : 1), inorganically precipitated CaCO3 should be either high-magnesium calcite (with 1.6–7.5 mol% Mg), or aragonite (due to the “Mg poisoning” effect which retards calcite precipitation). But, in fact, some low-magnesium calcite is precipitated in intracellular skeletons, because some organisms have succeeded in reducing Mg2+ diffusion through their cells during shell secretion (Veizer 1983). Increased pressure increases CaCO3 solubility: at 2 °C, calcite and aragonite are almost twice as soluble at 500 atm as at 1 atm. Carbonate sediments are rare below 4.5 km depth in the oceans. Aragonite dissolves at shallower depths than calcite, so pteropod oozes (aragonite) are generally shallower than foraminiferal oozes (calcite). Protective organic ﬁlms around grains and fecal pellets may retard the process until after burial, when porewaters may build up enough CaCO3 to preserve the sediment. 3.3.1.2 Marine carbonate sediments Marine carbonates are overwhelmingly biochemical sediments whose primary grain size, grain shape, composition, and sorting are based on the size, shape, taxonomy, and ecology of the preservable organisms living in an area. Biochemical carbonate sediments accumulate where skeletal production is undiluted by clastic input or unfavorable secreting or preservational environments. Thus, they tend to occur in areas isolated by distance or climate from land-derived material and in waters saturated in calcium carbonate. Deep oceanic calcareous oozes accumulate above the carbonate compensation depth while deep-shelf carbonates form on both tropical and temperate shelves in cool water. In the northern hemisphere, the present low sea level and freshwater input limits shallow-shelf carbonates to mainly tropical and subtropical areas, but in the southern hemisphere, around Australia and New Zealand, extensive shallow carbonate shelves also occur in temperate areas (James 1997). Carbonate mineralogy, chemistry, and isotopic ratios vary within the upper 1500 m of the oceans, so controls must also vary over this interval. Most carbonate sediments are organically precipitated, and the major ion

ratios of sea water are constant regardless of depth or latitude, so the dominant controls are temperature, ion concentration (salinity), and type of organism (Chave 1962). Temperature is the main control, since it partly controls ion concentration (salinity) and fundamentally controls the distribution of all life on earth. Ion concentration (salinity) is the dominant control on carbonate mineral saturation states and thus controls the possible mineralogy and chemistry of particles precipitated. Unfortunately, temperature and carbonate ion concentrations are strongly covariant, so their relative importance in controlling mineralogy cannot be assessed from available data (Burton & Walter 1987). Organisms control the actual mineralogy, chemistry, grain size, and structure of carbonate particles. In turn, the organisms present at any one time are controlled by a host of environmental factors, including light, temperature, salinity, and nature of the substrate. Furthermore, the type of organisms present is dependent on accidents of evolution and extinction. We therefore need to know: (i) the mineralogy, chemistry, and sizes of carbonate grains secreted by organisms and any variations due to environment; (ii) the environmental factors that control the distribution of carbonatesecreting organisms; and (iii) the effect of evolution and extinction on the variety and distribution of organisms though time. Most major taxa of organisms have a basic dominant mineralogy, although wide variations are possible (Table 3.3). Most also have variable non-equilibrium major elements and isotope ratios (Fig. 3.39). Obviously, the type and abundance of carbonate producers determines the initial mineralogy and chemistry of the sediment. Thus, tropical shallow-shelf areas dominated by green algae, corals, and molluscs have very different mineralogies and chemistries to tropical deep-shelf and temperate shallow-shelf areas dominated by red algae, bryozoa, and crinoids. The sizes of grains and the structures of skeletons vary with taxonomic position, ecology, and age, and these skeletons can break down into smaller constituent grains with disjunct size distributions (Fig. 3.41). Alternately, organisms and microorganisms can bind loose sediment together to form larger particles, such as ooids, pellets, and aggregates (Fig. 3.42). Light, temperature, salinity, and substrate (in order of decreasing scale) control aquatic communities. Within these, local factors such as oxygen content, food supply,

Sediments and sedimentary rocks 51

Fig. 3.41 Breakdown of green algae (Halimeda) and coral (Acropora) (from Flugel 1982, ﬁg. 30). (Copyright (1982). Reproduced with permission of Springer-Verlag.)

Fig. 3.42 Structures formed by binding loose carbonate sediment (from Tucker 1991, ﬁg. 4.1). (Reproduced with permission of Blackwell Publishing Ltd.)

turbulence, competition, variability, etc., produce a shifting mosaic of local communities (Fig. 3.43). Evolution and extinction have transformed the taxonomy and ecology of organisms. Paleozoic clear-water marine benthic communities are dominated by brachiopods, bryozoans, crinoids, corals, and trilobites;

mostly immobile, attached, and free-living suspension feeders that secrete calcite. Mesozoic to Recent clearwater marine benthic faunas are dominated by molluscs, corals, and ﬁsh; mostly mobile forms that secrete aragonite (Fig. 3.44).

Fig. 3.43 Various organic communities on a recent Dutch tidal ﬂat (from Reise 1985, ﬁg. 5.3). (Copyright (1985). Reproduced with permission of Springer-Verlag.)

Fig. 3.44 Reconstructions of benthic communities: (a) Devonian; (b) Jurassic (from McKerrow 1978, ﬁgs 37, 66). (Reproduced with permission of Alphabet and Image.)

Sediments and sedimentary rocks 53

Fig. 3.45 Changing taxonomy and mineralogy of marine organisms through time (from Brand & Morrison 1987, ﬁg. 7). (Reproduced with permission of the Geological Association of Canada.)

Thus, the mineralogies of biochemical limestones have changed through time (Fig. 3.45). Is this turnover a result of evolutionary, environmental, ecological, or accidental (catastrophic) change? Whatever the reasons, they make Paleozoic limestones difﬁcult to interpret. 3.3.1.3 Constituents Like clastics, carbonate constituents vary from boulders to clay size. Unlike clastics, the shapes of carbonate particles may be very peculiar. Skeletal grains Organisms secrete a bewildering variety of shapes, structures, and sizes of protective calcium carbonate skeletons, of a wide variety of compositions. Thus, ﬁne carbonate muds are precipitated by algae, all shapes of sand-sized grains are precipitated by foraminifera, spoon-like granule- and pebble-size grains are precipitated by various clams and brachiopods, and enormous massive, branching, and foliaceous boulderlike masses are precipitated by corals, algae, and other colonial organisms. These skeletons can then further

break down into a mixture of various shapes and sizes (Fig. 3.41), and can then be transported and reworked, which creates some special problems. First, the variety of skeletal grains initially present reﬂects the types of organisms present. These are controlled by local environmental factors that can only be inferred by knowing the tolerances and ecologies of the organisms themselves. However, skeletal grains from different environments can easily get mixed. Thus, the shape, size, and initial composition of the skeletal grains is not entirely the result of physical or chemical processes as in clastic and purely chemical sediments. Second, the skeletal grains of Recent sediments can be attributed to organisms whose environmental tolerances are known. This cannot be done directly for extinct organisms. Thus, the original control on carbonate composition may not be known. Third, the ﬁne sediment produced by the breakdown of skeletal grains cannot be identiﬁed with its precipitating organism. Since the geochemical and isotopic composition of a carbonate sediment is mostly dependent on the organism present, then any inferences about

54 Chapter 3 carbonate environments based on geochemistry or isotopic compositions have to be carefully evaluated. In fact, understanding most carbonate sediments involves, as a basis, knowing the taxonomy, ecology, and evolution of both fossil and recent organisms, the mineralogy and chemistry of their skeletons, and the sorts of particles they produce (see Flugel 1982; Scofﬁn 1987). Non-skeletal grains Non-skeletal grains vary from boulder-sized lumps eroded from preexisting limestone down to microcrystalline mud precipitated directly from solution. Clasts are sand-sized and large fragments of preexisting limestone deposited with contemporary carbonate sediments. Conventionally, such fragments are EXTRACLASTS (derived from outside the basin of deposition) or INTRACLASTS (derived from within the basin of deposition). It is often difﬁcult to tell them apart. Extraclasts are not common except where limestone outcrops undergo rapid physical weathering, for example in semi-arid environments and from cliffs. However, in some places, such as the coasts of Croatia, stream deposits may be composed entirely of extraclasts cemented by calcrete. Similar ancient limestones would be very difﬁcult to recognize and interpret. Intraclasts are limestone fragments nearly contemporary with the host sediment. They may be pieces of beachrock or hardgrounds or even bits of stromatolites. In contrast to extraclasts, intraclasts are common in limestone because many carbonate sediments either form as rock or lithify very early. Internal fabrics may be truncated showing that the intraclasts were lithiﬁed when transported. Aggregates form when several grains get stuck and cemented together, while coated grains form when carbonate layers are precipitated over single or aggregated grains. Both form where carbonate is precipitating but where some movement occurs that prevents entire layers being cemented. GRAPESTONES are knobbly aggregates of grains that resemble small bunches of grapes. The grains are often micritized and are weakly cemented by ﬁne-grained aragonite. Grapestones are found today in the sheltered shallow-water zones of the Bahamas and the Persian Gulf where they probably formed when the grains were at rest for some time in areas of strong supersaturation. Grapestones are not common outside the Bahamas and Persian Gulf and have rarely been described from ancient limestones.

OOIDS (or ooliths) are oval grains with onion-like internal laminae. They grow in soils, in caves, in hot springs, in lakes, and in the sea. Marine and non-marine ooids cannot be distinguished by their petrography or chemistry, so the simple presence of ooids is not sufﬁcient to deﬁne a marine environment. Ooid mineralogy and composition directly reﬂect the chemistry of the precipitating waters and the physico-chemical conditions of precipitation. Thus, soil ooids are mixtures of aragonite and high-Mg, high-Sr calcite, since they often form in hot, saline conditions in semi-arid soils. Cave ooids are usually calcite with low Mg and Sr, since they precipitate from cool fresh waters and are often derived from old stabilized limestones. Oncoids (and pisoids) are grains with irregular coatings usually precipitated by various organisms, often in moderately quiet water. PELOIDS are rounded grains of microcrystalline carbonate with an indeterminate origin. Many of them are the fecal pellets of organisms and they may show internal structures. Deposit-feeding animals that swallow ﬁne mud and produce abundant pellets include gastropods, worms, and shrimps. One pellet-producer may dominate an area and produce uniform sizes of pellets. Many peloids are small rounded intraclasts and micritized skeletal grains. So, unless the pellets can be deﬁnitely recognized as fecal, the general term peloid is appropriate. Carbonate muds are less than 64 mm (4 phi) in size, and form in very diverse ways. Single aragonite and Mg-calcite needles, 1–10 mm in length, form most of the clay-sized fraction over vast areas of warm, shallow banks. Such crystals can be produced both by inorganic precipitation experiments and by natural disintegration of some calcareous green algae. Shinn et al. (1989) observed that milky clouds of such crystals formed in open water, aggregated into silt- and sand-sized ﬂoccules, and settled to the bottom.

3.3.1.4 Transportation and deposition Once formed, carbonate particles are abraded, sorted, rounded, and eventually destroyed by the whole spectrum of agents of transportation and deposition. As with clastic sediments, the grain size, shape, sorting, and roundness of bioclastic carbonate sediments can be used to infer the nature of transportation and deposition. However, the enormous variety of skeletal mineralogies, structures, shapes, and densities of organisms can undergo very selective sorting and

Sediments and sedimentary rocks 55 100

Piaster 6 gm

Corallina 7 gm

ne

Echinoids 8 gm

Limpets 6 gm

Tegula 9 gm

Haliotis 17 gm

20

Aletes 23 gm

Original material (%)

60

Mytilus 48 gm

Go

M.

A.

H.

T.

L. E.

0

P.

C.

2h

0 time 100

Go

Go

ne

ne

80 M.

A.

H.

T.

L.

20

M.

A.

H.

0

T.

L.

183 h

40 h

Fig. 3.46 Differential abrasion and loss of shell types with transportation: results of a tumbling-barrel experiment with sand and shells greater than 2 mm (from Chave 1964, ﬁg. 3).

destruction. What survives to form sediment may be very different in size, shape, composition, and proportion to what was there before movement (Fig. 3.46). 3.3.1.5 Classiﬁcation Carbonate sediments are moved around like clastic sediments, so a simple classiﬁcation based on grain size can be used for clastic limestones (Fig. 3.38). However, this is inadequate for carbonates that have not been moved much, and it shows neither the constituents nor the textures of the limestone (though modiﬁers can be added, e.g. crinoidal calcarenite). Two classiﬁcations based on constituents and textures can accurately and adequately describe most carbonates. Folk’s classiﬁcation emphasizes constituents (Fig. 3.47), and is useful when comparing carbonates from different Recent and ancient environments. Dunham’s classiﬁcation emphasizes depositional texture (Fig. 3.48 and Table 3.4), and is useful when comparing carbonates deposited in different ways.

Distinguishing grainstones from packstones can be a problem, as carbonate mud matrixes are liable to recrystallize to coarser spar. However, the two-stage cements of grainstones contrast with the variable grain size of recrystallized mud (Fig. 3.49). Another problem is how the packing of shells may be misleadingly open, giving a ﬂoating wackestone appearance to a packstone or grainstone framework (Fig. 3.50). Combining Folk and Dunham terms usefully describes both the constituents and the texture of limestones; for example, biosparite grainstone. The preﬁx dolo- can be used for limestones replaced by dolomite. Different combinations of allochems and matrix are deposited by the varying processes of mass ﬂows, water ﬂows, and wind in different marine environments. 3.3.1.6 Diagenesis Carbonate diagenesis is basically the transformation of aragonitic and calcitic sediment into limestone or dolomite. Dry transformations of aragonite to calcite need temperatures of about 400 °C, take a long time,

56 Chapter 3

Fig. 3.47 Folks’ constituent classiﬁcation (from Tucker 1991, ﬁg. 4.35). (Reproduced with permission of Blackwell Publishing Ltd.)

Fig. 3.48 Dunham’s depositional texture classiﬁcation (from Greensmith 1978, ﬁg. 8.4). (Reproduced with permission of HarperCollins.)

and mainly interest metamorphic petrologists; wet transformations of aragonite to calcite (in water lacking free Mg2+) are much faster, taking only a few days. Primary calcite sediments undergo a different diagenesis

regardless of any other factors, which is why diagenetic fabrics of Paleozoic calcite limestones resemble those of Recent temperate calcite limestones. Carbonate diagenesis includes cementation,

Sediments and sedimentary rocks 57 Table 3.4 Boundaries of Dunham’s classiﬁcation (from Greensmith 1978, table 8.2). (Reproduced with permission of HarperCollins.) Depositional texture recognizable Contains mud Mud-supported less than 10% grains Mudstone

more than 10% grains Wackestone

Not

Grainsupported

Lacks mud and is grainsupported

Original components bound together

Packstone

Grainstone

Boundstone

recognizable

Crystalline carbonate

Fig. 3.49 (a) Two-stage spar cement of grainstone; (b) variable grain size of recrystallized mud (both modiﬁed from Flugel 1982, ﬁg. 8A,B). (Copyright (1982). Reproduced with permission of Springer-Verlag.)

Fig. 3.50 (a) Open shell framework; (b) compact ooid framework (from Blatt 1982, ﬁg. 1.21).

dissolution, neomorphism, grain compaction, lithiﬁcation, and replacement, with all the loss of information these entail, and it begins as soon as the sediment gets buried. Cementation Cementation, like precipitation, is controlled by near-surface conditions, so the mineralogy and isotopic composition of early cements tend to be similar to those of the parent sediment. In warm,

normal sea water (e.g. in Recent shallow tropical areas), nuclei and impurities abound, CaCO3 is highly supersaturated, and precipitation is rapid. Early diagenetic cements thus tend to consist of rapidly precipitated ﬁne-grained crystals. In cold, normal sea water (e.g. in Recent polar, temperate, and deeper water areas), CaCO3 saturation and precipitation rates are lower and early diagenetic cements should be coarser grained. Aragonite and high-Mg calcite dominate early

58 Chapter 3 warm-water cements while low-Mg calcite dominates early cold-water cements. For example, in modern seas, high-Mg calcite (12 mol% MgCO3) dominates at depths between 100 and 1000 m while low-Mg calcite (3–5 mol% MgCO3) dominates below about 1000 m. Intergranular cementation, sufﬁcient to harden sediment into a rock, may start a few centimeters below the sea ﬂoor but requires the following conditions: 1 supersaturated porewaters (warm water at present); 2 a high water ﬂux through the sediments, otherwise CaCO3 gets depleted at the site of cementation (as sea water contains only about 0.05 g CaCO3 per liter, tens of thousands of pore volumes must pass through a pore to ﬁll it with calcium carbonate, which can normally occur only in agitated water or over a very long time); 3 porous sediments to allow such a ﬂux (e.g. grainstones); 4 a stable substrate, so that the grains are exposed to ﬂushing for a long time, yet newly formed delicate cements are not abraded. These conditions usually occur only in speciﬁc areas or within speciﬁc layers. Frequently, waves and currents remove soft sediment down to a cemented layer whose exposed surface or HARDGROUND can then be colonized by organisms. Hardgrounds typically have uneven, irregular bottom surfaces and smooth or rolling top surfaces, which may, if exposed for a long time, become impregnated with iron, manganese, or phosphate. Recent active intergranular cementation is recorded from shallow-water bioherms, deeper platform margins and seamount tops, the stationary lobes of ooid bars, and below tropical beaches. Irregular nodular cementation is more common in areas of sluggish sea-water circulation where carbonate may segregate around shell fragments or diffuse toward centers of nucleation. Dissolution Dissolution occurs where undersaturated waters contact the sediment or rock in freshwater and Recent shallow temperate, polar and deep seas, and where pressure solution takes place during burial. Pressure solution gives CaCO3-rich porewaters, which can migrate and further cement adjacent rocks, and forms stylolites which can indicate the amount of carbonate lost. Aragonite can preferentially dissolve leaving moulds and increasing secondary porosity. Neomorphism Neomorphism involves replacement by material of the same chemical composition. Degrading neomorphism, or MICRITIZATION, occurs where crystals are made smaller by replacement. This normally occurs

from the outside inwards, destroying the original crystalline texture and giving the grains a chalky appearance. Much of this micritization is caused by boring algae and bacteria. Such grains are particularly abundant in shallow lagoons where algal mats stabilize the surface and where cyanobacteria often coat interstitial grains. Aggrading neomorphism occurs where aragonite and high-Mg calcite needles change to coarser, blocky low-Mg calcite crystals. Grain compaction Grain compaction increases grain packing, decreases porosity (and often permeability), and can culminate in pressure solution welding the sediment into a hard massive rock. Lithiﬁcation Lithiﬁcation transforms the uncemented or cemented sediment into a hard rock. It involves ﬁlling in any residual pores and obviously overlaps with cementation. Replacement Replacement changes the chemical composition of the carbonate. Common replacements are dolomite and silica (chert). The dolomite replacement problem has been around for a long time. Dolomite replacement may be very early (e.g. some laminae are replaced while the underlying and overlying carbonate remains unaltered), very late (e.g. when the entire rock is replaced by dolomite), or sometimes in between (Fig. 3.51). Migrating Mg-enriched solutions (formed by the loss of magnesium from metastable high-Mg calcite) may replace entire limestones, forming secondary dolomite. However, the amount of Mg needed is very large. To convert a typical calcium carbonate sediment, containing 6% MgCO3 with 40% porosity, to dolomite requires 807 pore volumes of sea water to ﬂow through the rock. With sea water diluted ten times with fresh water in a mixing zone, then about 8100 pore volumes are needed. With halite-saturated brine, only 44 pore volumes are needed. But even with these large volumes, porosity is not decreased. To reduce porosities to values common in ancient dolomites requires even greater volumes. Since smaller volumes of saline water are needed to dolomitize limestone, and since most modern dolomites are associated with evaporitic conditions, then dolomite is mostly an evaporite mineral. Geochemists have attempted to mimic the surface conditions of modern dolomite formation in the laboratory and have failed. However, a dog has succeeded: at 38 °C and one atmospheric pressure, a Dalmatian managed to precipitate dolomite uroliths in

Sediments and sedimentary rocks 59

Fig. 3.51 (a) Late dolomite mosaic; (b) early scattered dolomite rhombs in micrite; (c) early synsedimentary dolomite laminae (small crystals) in micrite. Fields of view are about 2 mm (from Williams et al. 1954, ﬁgs 121A, 119A,B). (© 1954 by W. H. Freemen. Used with permission.)

its urinary tract, where the physical and chemical conditions may be more like natural systems than during laboratory experiments on dolomite formation (Warren 1989). All these diagenetic changes alter the character of the original sediment and have to be determined and allowed for before the depositional processes can be worked out. 3.3.2 Cherts Cherts are nodules, layers, or beds of colloidal silica and they are normally diagenetic (Fig. 3.52a). Exceptions are the deep-sea radiolarian and shallow-water diatom cherts (precipitated by micro-organisms) and hot-spring cherts (precipitated from hot waters) (Fig. 3.52b). Secondary cherts form by replacement of soils (e.g. silcretes), carbonates (e.g. chert nodules in limestone), and organisms (e.g. siliciﬁed wood), and may be remobilized to precipitate in cracks (Fig. 3.52b). Interesting precipitates are the alternating laminae of ferric oxide and chert in Archean banded ironstones. Most are between 2400 and 1800 million years old and pre-date an oxidizing atmosphere. The alternating laminae may be due to pH and Eh changes, either during seasonal ﬂuctuations or during the removal of CO2 and production of O2 by photosynthetic cyanobacteria

during the day (oxidation of iron) and the respiratory production of CO2 at night (precipitation of silica). 3.3.3 Evaporites EVAPORITES (sulphates and chlorides) form where evapora-

tion exceeds precipitation. Although this need not imply warm conditions (since there are saline lakes in Antarctica), thick evaporite deposits form only in large, arid, semi-enclosed and continental sedimentary basins. Since chemical rocks readily recrystallize and redissolve, the only thing you can normally tell from them is the chemical state of the precipitating solution (Warren 1989). Evaporites form in three main ways: (i) by evaporation of sea water in isolated seas; (ii) by evaporation of fresh water in isolated lakes; and (iii) by evaporation of temporary pans and groundwaters in sediments. 1 Most evaporites form by sea water evaporating in semi-enclosed basins. At present, salinity in the open ocean is 34–36‰, rising to 40‰ in semi-enclosed areas furthest from the ocean. The evaporation of normal sea water produces a characteristic sequence of salts (Table 3.5). The ﬁrst compound to precipitate is calcium carbonate (CaCO3, usually as aragonite) at twice normal salinity (40–60‰). This is followed by calcium sulphate (CaSO4.2H2O, gypsum) at ﬁve times normal salinity (130–160‰); sodium chloride (NaCl, halite) at 11–12

60 Chapter 3

Fig. 3.52 (a) Dolomite rhombs in a replacement chert matrix; (b) primary radiolarian chert. Fields of view are 2.5 mm and 1 mm respectively (from Williams et al. 1954, ﬁgs 123, 124). (© 1954 by W. H. Freeman. Used with permission.)

Table 3.5 Mineral sequence from sea-water evaporation (from Warren 1989, table 1.2). Mineral

Concentration factor

% Water loss

Brine density

K-Mg salts Halite Gypsum (anhydrite) CaCO3 Sea water

63 ¥ 11 ¥ 5¥ 2–3 ¥ 1¥

98.7 90 80 50 0

1.29 1.214 1.126 1.10 1.04

times normal salinity (340–360‰); and various bittern salts of potassium and magnesium at concentrations of more than 60 times normal salinity. Large Recent marine evaporite basins on the scale of ancient ones do not exist (but see Logan 1987). Ancient basins possibly accumulated salts in three main ways (Fig. 3.53). The salinities of these partially enclosed seas were apparently very stable, staying at halite- and gypsumprecipitating concentrations for long periods of time. How deep the water actually was during precipitation is debatable, even in the examples given (Kendall 1992). 2 Lake evaporites have peculiar minerals that are absent from marine evaporites because continental water chemistry is far more variable then sea water. More diverse evaporite mineral assemblages and sequences are the result. Regardless of the water type, the ﬁrst miner-

als to precipitate are the alkaline earth carbonates, low-Mg calcite, high-Mg calcite, aragonite and dolomite, with gypsum usually next. Subsequent precipitates depend on the brine inﬂow chemistry. 3 Subaerial precipitation of gypsum and anhydrite occurs together with dolomite in tidal ﬂats and very shallow lake pans. Sabkha evaporites form in supratidal areas, typically passing seaward into limestone and landward into red continental shales and sandstones. Pan evaporites are deposited in shallow salt pans that lie above normal high tide and ﬂood only during spring tides or storms. Between ﬂoods, the pans evaporate to the point of gypsum precipitation; only rarely does halite form. Removal of calcium from pan water causes the magnesium/calcium ratio to increase, promoting dolomite replacement.

Sediments and sedimentary rocks 61

Fig. 3.53 Three mechanisms for evaporite formation (from Prothero & Schwab 1996, ﬁg. 14.9).

62 Chapter 3

Fig. 3.54 Raised swamp model for thick peats and coals (from McCabe 1984, ﬁg. 12). (Reproduced with permission of Blackwell Publishing Ltd.)

Diagenesis of soluble evaporites can be complicated. When buried gypsum rises above 60 °C it is transformed into anhydrite, releasing calcium sulphate saturated water of dehydration, increasing the porosity of the sediment by 38%, and decreasing the strength of the anhydrite due to increased lubrication. If the water cannot drain freely then the sediment may liquify, causing intense synsedimentary deformation. If the water escapes, then the differential sediment compaction causes deformation of the overlying beds (Warren 1989). Salt deforms even more easily and on a larger scale, forming stratigraphic and structural traps for oil and gas. 3.3.4 Peat and coal Peat and coal are the ultimate biochemical sediments and rocks, since the properties of the peat (and eventually coal) depend mostly on the plant community (controlled by the environment) and the type of plant present (controlled by evolution) (Moore 1995). PEAT is a mixture of partially decomposed plant remains, and can form anywhere plants grow and accumulate. Diagenesis, involving physical and chemical changes, then transforms peat through a number of different coals to anthracite. However, since most organic matter is rapidly oxidized in soils and recycled back into plants, peats only form if decomposition is slowed down, and this happens only if the organic matter accumulates in relatively stagnant and oxygen-deﬁcient water. To form peat, therefore, the water table must be in equilibrium at or above the sediment surface. If the water table rises faster than the plants accumulate, then the swamp will drown. If the water table drops, then the swamp will dry out and the peat will oxidize and erode. Peat production is a balance between plant production and decay, which are both controlled by climate (McCabe 1984).

Most present-day peats form in cool climates between 50°N and 70°N, and high rainfall and temperature are not necessary. Thus, the vast peat swamps of Canada and Siberia form in cool areas with rainfall of less than 50 cm/year (Martini & Glooschenko 1985). Most tropical rain forests have no peat since organic matter is rapidly recycled through the thin soils. Tropical peats now form only in parts of Southeast Asia where over 200 cm of rainfall occurs throughout the year. Peat (and hence coal) is thus not a good paleoclimate indicator. For thick peats to form, plant remains have to accumulate under anaerobic conditions, and the inﬂow of oxidizing water and clastic sediments has to be limited. This can happen in three ways: 1 Lake basins can become isolated from river and clastic input and develop ﬂoating swamps. However, any tectonic, climatic, river-course, or lake-level changes will tend to change the environment. Floating swamps are probably too ephemeral to form thick peat deposits. 2 Low-lying areas can get cut off from water inﬂow and clastic input; for example locally in meander cut-offs and more extensively in large ﬂoodplain swamps. However, for thick peats (and coals) to form in such areas, they must be isolated for long periods of time. Although these swamp types are often used to explain peat and coal deposition, they are also sinks for any clastic sediments. Flooding river channels tend to deposit sandy crevasse-splay deposits in the peat. 3 Vegetation can accumulate and raise swamps above river, lake, and sea levels so that they are protected from water inﬂow and clastic input, as in the raised swamps of Southeast Asia. The vertical growth of the swamp conﬁnes river channels, resulting in stacked anastomosing channel systems separated by thick peats (Fig. 3.54). Although raised swamps tend to dehydrate, oxidize, and erode during dry spells, and when they rise above

Sediments and sedimentary rocks 63

Fig. 3.55 Floras and the resulting coal: comparisons between (a) Carboniferous and (b) Miocene swamps (from Teichmuller 1989, ﬁgs 8, 13). (Copyright (1989), with permission from Elsevier Science.)

groundwater levels, they do provide a site where thick undiluted plant remains can accumulate and are a good model for thick ancient coals. Swamps can evolve from ﬂoating swamps through low-lying and raised swamps to dehydrated swamps. The original ﬂoral and peat constituents are highly modiﬁed during the chemical and physical changes that transform peat to coal. Nevertheless, the coal constituents and their distribution within the coal can help you work out how and where it formed (Bustin et al. 1985). The main constituents of coal (macerals) are vitrinite/huminite, derived from woody tissue and bark; exinite, derived from spores, pollen, resin, cuticle, and algae; and inertinite, derived from various constituents by the early loss of volatiles by oxidation, charring, fungal attack, etc. These, with variable amounts of mineral matter, form four main coal lithotypes: (i) vitrain has very bright thin bands or lenses of vitrinite; (ii) durain has exinite-dominated dull bands;

(iii) fusain has black ﬁbrous bands dominated by inertinite: and (iv) clarain has ﬁne, alternate dull and bright laminae, and is a mixture of vitrinite, exinite, and intertinite. The distribution of these lithotypes reﬂects the original plant composition and how this was transformed into coal. However, because the ﬂoras that form peats and eventually coals have evolved through time, it can be difﬁcult to interpret how ancient coals formed (Fig. 3.55). Late Cretaceous and Tertiary ﬂoras are taxonomically close enough to compare with modern ﬂoras, although early Tertiary ﬂoras lack a savannah or grassland component (Shearer et al. 1995). Late Paleozoic ﬂoras are very different: no ﬂowering plants, fewer gymnosperms, and ﬂoras dominated by ferns and extinct seed-ferns (Phillips 1981). The ﬂoras of a coal seam can, to a certain extent, be reconstructed from its constituents, and the vertical changes in constituents may give some clues about environmental and vegetation changes during the development of the swamps.

64 Chapter 3

Fig. 3.56 (a) Limonite ooids in shelly limestone; (b) chamosite ooids in sandstone. Both are Lower Jurassic ironstones, UK. Fields of view are about 2 mm (from Williams et al. 1954, ﬁg. 127). (© 1954 by W. H. Freeman. Used with permission.)

Fig. 3.57 Stratigraphy of Permian phosphorites in Idaho and Wyoming (from McKelvey et al. 1959).

Sediments and sedimentary rocks 65 3.3.5 Ironstones and phosphates Ironstones and phosphates are volumetrically insigniﬁcant yet economically important. Both occur with clastic sediments but need reduced rates of sedimentation to form. IRONSTONES may consist of sulphides (pyrite), carbonates (siderite), silicates (chamosite and glauconite), and oxides (goethite and hematite). Most, with the exception of oxides, form experimentally under reducing conditions. Pyrite and siderite form naturally by diagenetic segregation and replacement in impermeable sediments. Siderite often forms nodules and band shales. Chamosite (berthierine) and glauconite form by replacement of clay minerals under reducing conditions in shallow- and deep-shelf waters, respectively. Goethite and hematite are usually secondary oxidation products of the other minerals in soils and sediments. There are no modern equivalents of the Phanerozoic chamosite oolitic iron ores, which often get oxidized to limonite (Fig. 3.56), or the Precambrian banded hematite/chert

iron ores. Because both primary ironstone types are typically interbedded with clastics, the main factor restricting their formation is concentrating enough iron while excluding clastics. The amount of dissolved or colloidal iron is very small in both sea water (< 0.01 ppm) and fresh water (< 0.1 ppm), so unusual conditions are needed to explain the primary oolitic ironstones and iron formations of the past (Bayer 1989). PHOSPHATES consist of various calcium phosphate halides and hydroxides (ﬂuorapatites, chlorapatite, hydroxyapatite). These form by accretion as nodules in the deep sea and as replacements of other minerals in shelf and slope environments (Slansky 1986). In shelf and slope environments, phosphates form where upwelling waters cause phytoplankton blooms which extract phosphate and concentrate it in condensed bottom sediments when they die. Most ancient phosphates probably formed near upwellings in areas with reduced rates of sedimentation (Fig. 3.57).

4 Major environmental complexes and their recognition 4.1 4.2 4.3 4.4 4.5 4.6

Introduction Impact and volcanic environments Continental environments Environments under water Mixed environments Peculiar environments

4.1 Introduction An environment is a part of the earth that is physically, chemically, and biologically distinct from another part of the earth. Basically, environments are deﬁned on aggregates of processes, and their boundaries can be sharp and distinct or gradational and blurred. Variation and change within an environment may be caused either by internal or external factors. Internal factors are part of the system, such as the lateral migration of river channels and the seasonal control of vegetation and ﬂooding. Stratigraphic sections can show repetitive, consistent, and gradual changes, such as repetitive ﬁning-upwards cycles of superimposed migrating river channels. External factors are outside the system, such as changes in climate and tectonics. For example, stratigraphic sections of meandering stream deposits may show unusual, random, and abrupt changes, such as interbeds of wind-blown sand (perhaps suggest-

ing a drier climate) and coarse conglomerates (perhaps suggesting tectonic uplift). However, such changes could also be caused by internal factors, or by combinations of internal and external factors. The wind-blown sand might reﬂect local migration of dunes, a drier climate, or changes in wind direction; the coarse conglomerates might reﬂect switching fan channels, a wetter climate, or tectonic uplift. Students should research at least one environment in detail for themselves, so they can appreciate the problems and uncertainties in even supposed straightforward recent–ancient comparisons. How you deﬁne environments, at what scale and with what emphasis, depends on the type of study. A geomorphologist may simply divide a river basin into braided or meandering channel and overbank environments; or he may further divide a meander channel into distinct subenvironments such as levee, point bar, oxbow lake, swamp, and so on. A sedimentologist may describe the

Major environmental complexes and their recognition 67 sediments of these units and the processes that form them. A geochemist may compare the chemistry of distinct lithologies, or adjacent pore spaces. A biologist (or paleontologist) may use organisms to distinguish overbank swamp and lake clays; or he may subdivide still further into shallow and deep lake clays, each with its own characteristic biota. There are numerous good textbooks on the nature and recognition of various environments: Davis (1992), Einsele (2000), Prothero and Schwab (1996), Reading (1996), Selley (1985, 2000), and Walker and James (1992) are good general texts. See Allen (1985), Chamley (1989), Fraser (1989), Miall (1996) for clastics; Flugel (1982), Scholle et al. 1983), and Tucker and Wright (1990) for carbonates; Hardie (1984), Melvin (1991), and Warren (1989) for chemical sediments; and Melosh (1989) and Rampino (1994) for impact deposits. Recognizing environments in ancient rocks requires you to infer the range of processes (physical, chemical, and biological) that characterize these environments from criteria obtainable from the rocks themselves (Chapter 3). Identifying ancient environments requires knowledge of the type, intensity, variability, and persistence of processes characteristic of each environment. Environmental divisions are end-members or extremes in a continuum and many intermediates and alternations can occur (Table 4.1 and Fig. 4.1). Nevertheless, these extremes do show the possible combinations of processes that occur in the commonest environments. A useful procedure is to deﬁne the ARCHITECTURAL ELEMENTS of a system; that is, the basic minimum of types of sedimentary units you need to adequately deﬁne the process variation in the system (Miall 1990). These elements can then be combined in various ways to describe speciﬁc environments, in the same way that different architectural components can be used to describe a building (Fig. 4.2). Modern environments should be used for comparison ﬁrst, before comparisons are made among supposedly equivalent ancient environments. Using conditions inferred from one ancient environment to interpret another simply leads to circular reasoning.

4.2 Impact and volcanic environments Impact and volcanic environments are distinctive because much of the material is thrown up by explosions and simply settles by gravity onto the earth. However,

the material may then be reworked by all other agents of transportation and deposition, such as gravity, water, air, and wind. 4.2.1 Impact environments Impact environments were only recently recognized on earth. They form layers within other environments and are often difﬁcult to identify. Near the impact site, deposits resemble mass ﬂow deposits from exploding volcanoes or collapsing slopes, but may contain high-pressure effects and impact melt rocks (Fig. 4.3). Further away, thin impact layers may be very difﬁcult to distinguish from volcanic ashes without detailed geochemical analysis (Alvarez et al. 1980). 4.2.2 Volcanic environments Volcanic environments vary from those dominated by gentle eruptions to those dominated by explosive eruptions (Cas & Wright 1987; Fischer & Schmincke 1984). Gentle basaltic eruptions build up easily identiﬁable lava successions, while explosive acidic eruptions deposit mostly volcaniclastics. Intermediate andesitic eruptions tend to be explosive but with more lava ﬂows. In all, lavas, mass ﬂow deposits and suspension fallout deposits interﬁnger with water-, wind-, and even ice-reworked material (Fig. 4.4). Volcanic environments are usually difﬁcult to correlate and date accurately because: (i) the volcanic sediments, lavas, and associated intrusions vary a lot over short distances, both in character and thickness; and (ii) they rarely contain any material suitable for dating (Orton 1996). Volcanic deposits on other planets resemble those on earth, though modiﬁcations related to gravity, density, and viscosity have to be made in order to understand them.

4.3 Continental environments Environments on land are distinctive because transport and deposition are dominated by the ﬂow of immiscible denser ﬂuids (mass ﬂows, water, and ice) beneath an overlying ﬂuid (air). This causes sharp contacts between the various layers and great variation in their properties. Continental depositional environments are dominated by clastic sediments. Biochemical and chemical sediments are found mostly in temporarily or permanently wet areas: peat in swamps, carbonates and chemical sediments in soils and lakes.

68 Chapter 4 Table 4.1 Simpliﬁed classiﬁcation of depositional environments. Main depositional process

Major environment

Mass ﬂow

Impact Volcanic

Fluid ﬂow Subaerial

Basaltic Intermediate Acidic

Ice-dominated

Ice-contact Braid plain Glacial lake

Water-dominated

Alluvial fan Braided stream Meandering stream Incised stream Wadi fan Ephemeral stream Saline lake and playa Sand sea Loess

Wind-dominated

Subaqueous

Subdivisions

Clastic system Shoreline River-dominated Tide-dominated Wave-dominated Continental shelf Continental slope Deep sea Carbonate systems Shoreline Shelf

Delta Estuary Tidal ﬂat Barrier bar/lagoon Beach

Deep-sea fan Abyssal plain

Ramp Rimmed shelf Reef

Slope Deep sea Epicontinental sea

4.3.1 Ice-dominated systems Ice-dominated systems are fed by glaciers. They occur in cold areas dominated by physical weathering. In tropical mountains, glaciers may extend into lower, warmer, and drier areas. Ice-dominated systems contain strata deposited by various combinations of ice ﬂows, mass ﬂows, water ﬂows, and wind. The predominant agents, however, are ice and cold-water ﬂows, resulting in deposits dominated by immature clastics (Fig. 4.5).

Chemical deposits are rare and biological systems unimportant. There are three main glacial sub-environments: icecontact deposits, braid plains, and lakes (it is convenient to include marine glacial deposits within the lake subenvironment here). 1 ICE-CONTACT DEPOSITS form next to a glacier and contain material deposited directly from melting ice (Fig. 4.6). 2 Braid-plain gravel and sand deposits form from melt-

Major environmental complexes and their recognition 69

Fig. 4.1 Major clastic environments and simple sediment characterization.

water streams and, apart from their mineralogical immaturity, are difﬁcult to distinguish from other braidedstream deposits without associated ice-contact deposits or evidence of freeze–thaw cycles (Section 4.3.2). 3 Glacial lake deposits resemble those of non-glacial lakes but with speciﬁc features caused by annual freezing and thawing and subaqueous glacier input. Marine glacial deposits have the same characteristics. At spring break-up, melting winter ice may ﬂoat away from the shoreline and drop enclosed large beach pebbles into offshore clays. Varved clays may form as annual winter freezing causes organic-rich clay particles to settle onto ﬁne silt and sand layers deposited from spring and summer meltwater input. Coarse subaqeuous ice-contact delta deposits may interﬁnger with these ﬁne lake or marine sediments.

4.3.2 Water-dominated systems Water-dominated systems are fed by perennial streams, which occur anywhere but mostly in wet temperate and tropical areas. Water-dominated systems produce strata that were deposited by combinations of mass ﬂows and water ﬂow (Fig. 4.7). The dominant process of persistent water ﬂow forms relatively mature, well-sorted clastics except near the source. Chemical deposits are rare, but biological systems are very important. Vegetation stabilizes the sediment surface, increases chemical weathering, and greatly modiﬁes the soil proﬁle. River systems change with decreasing rainfall into semi-arid and desert systems. With decreasing competence downstream, streams in humid climates change from alluvial fans through braid plains, meandering alluvial plains,

70 Chapter 4

Fig. 4.2 Architectural elements: (a) tabular river sand bodies on a large scale; (b) lensitic channel sand bodies seen in part (a); (c) individual bedforms that make up each channel deposit (from Miall 1990, ﬁg. 4.56). (Copyright (1990). Reproduced with permission of Springer-Verlag.)

and incised (anastomosing) alluvial plains before they enter a lake or the sea. Because of their great variety, stream deposits are best described with individual architectural elements that summarize the great variability in processes (Miall 1992).

which builds out from the Himalaya into the Ganges valley (Fig. 4.8). The low slopes on such fans make them difﬁcult to distinguish from braid and meandering plain systems in ancient deposits. 4.3.2.2 Braid plains

4.3.2.1 Alluvial fans ALLUVIAL FANS are formed where streams leave mountain

gorges and dump their loads in adjacent valleys. Streamdominated or “humid” fans show a gradual reduction in slope compared to fans in arid regions (Section 4.3.3). The largest described humid fan is the Kosi River fan,

BRAID PLAINS are formed by braided streams, which have multiple channels caused by highly ﬂuctuating sediment-laden discharges alternately ﬂushing out and blocking temporary channels (Fig. 4.9). Rapid deposition of coarser material in various types of bars blocks unstable channels which then spread out across the

Major environmental complexes and their recognition 71

Fig. 4.3 Sedimentary sequences around an impact into shallow sea: E, ejecta (Ea, subaerial; Eh, water-deposited; Er, reworked); D, disturbed material (Dd, dropstone; Dc, compression; Db, brecciated); M, mobilized material (Mt, turbidity current; Ms, slumps; Mb, backwash; Mr, reworked); d, dropstone; m, megaclasts; i, injected material; l, lamination; b, slump block; c, cross-bedding; f, reworked trees. Column 1 is an oceanic section a long way from the impact. The dashed line is the original surface (from Rampino 1994, ﬁg. 7). (Reproduced with permission of the University of Chicago Press.)

72 Chapter 4

Fig. 4.4 Characteristics of acidic volcanic environments around an island arc volcano (from Einsele 1992, ﬁgs 2.23b, 2.24e, 2.26b). (Copyright (1992). Reproduced with permission of Springer-Verlag.)

ﬂoodplain into a multiple braided pattern. Fine sediment is normally carried downstream except where small ponds develop on the braid plain after peak discharges. Gravel braid plains pass downstream into sand braid plains, which themselves grade into meandering stream

plains. Sections tend to consist of irregular cut-andﬁll, ﬁning-upwards cross-bedded gravels and sands. In humid climates braid plains differ little from alluvial fans.

Major environmental complexes and their recognition 73

Fig. 4.5 Types of glaciers, glacial environments and landforms (from Edwards 1986, ﬁg. 13.2). (Reproduced with permission of Blackwell Publishing Ltd.)

4.3.2.3 Alluvial plains plains are formed by meandering and incised streams, and develop where gradients and discharge ﬂuctuations are smaller than for braided streams. Meandering streams show an organized separation of channel and overbank deposits (Fig. 4.10). The meandering channel migrates laterally, depositing characteristic ﬁning-upwards tabular channel sand bodies within ﬁner overbank deposits. Each sand body consists of a coarse basal lag overlain by unidirectional, but scattered, large-scale cross-bedded point-bar sands, which in turn are overlain by smaller-scale festoon cross-bedded chute-and-pool sands. Floods sporadically break though the channel levees forming crevasse-splay deposits on the alluvial plain. Meander cut-offs may become ﬁlled in with ﬁne organic-rich lake and swamp deposits. Overbank deposits of ﬁne sand, silt, and clay settle out of suspension when a river breaks its banks and inundates the ﬂoodplain. After drying out, the ﬂoodplain may develop immature soils, including caliche in seasonal climates. Low-lying areas near channels may stay

ALLUVIAL

waterlogged, accumulate organic matter, and develop peats. INCISED STREAMS have anastomosing channels stabilized by vegetation, and normally develop in areas with very shallow slopes such as marshy and swampy areas. Sedimentation is normally by vertical accretion, with stacked shoestring channel sands passing laterally into alternating marsh and swamp clays and peats/coals. If sand is deposited in channels at a faster rate than dead vegetation accumulates in marshes and swamps, then channels with well-developed levees may rise above the adjacent swamp and marsh deposits. Channels then change by crevassing followed by rapid avulsion, and deposits are random arrays of shoestring channel sands bounded by ﬁner overbank sediments with occasional crevasse-splay deposits (Fig. 4.11). If dead vegetation accumulates in marshes and swamps faster than sand deposition in channels, then thick raised peats may build up. These conﬁne the channels in the same place for long periods of time and the resulting deposits consist of stacked channel sands separated by thick peats and clays. Deposits of this type

74 Chapter 4

Fig. 4.6 Ice-contact features and their deposits (modiﬁed from Edwards 1986, ﬁg. 13.11). (Reproduced with permission of Blackwell Publishing Ltd.)

form the very thick coal seams of the Pennsylvanian of eastern North America and the Eocene of western North America. Normally, of course, conditions ﬂuctuate and alternate with periods of incision due to climatic, tectonic, or water-level changes. In all these river environments, small interchannel lakes do not last long, as ﬁne sediment and vegetation rapidly builds up and soon transforms them into swamps or marshes. Large lakes are more persistent. 4.3.2.4 Large lakes Large lakes form in structural depressions: they are very diverse but temporary and are eventually ﬁlled in with

sediment. Perennial streams supply lakes in water-dominated systems, which thus tend to be hydrologically open and do not precipitate much biochemical or chemical sediment. River input, wave reworking, and turbidity currents dominate the processes in such open lakes (Fig. 4.12). Simple deltas have steep foresets overlain by ﬂat topsets. Gravel and sand beaches, sand spits, and small barriers form in wave- and current-dominated areas. Offshore, mass ﬂows, turbidity currents, and pelagic fall-out form pebbly mudstones, graded sandstone, organic-rich clays, and minor biochemical layers. Cyanobacteria may form stromatolites and oncoids in shallow water, associated with shells of gastropods, bivalves, and ostracods (Fig. 4.13).

Major environmental complexes and their recognition 75

Fig. 4.7 Wet alluvial systems (from Einsele 1992, ﬁg. 2.8). (Copyright (1992). Reproduced with permission of Springer-Verlag.)

Ancient, large, open-lake deposits may be difﬁcult to distinguish from marine deposits but should show many more water-level ﬂuctuations. In Phanerozic times, marine fossils should be absent, but a low-diversity fauna

including gastropods, bivalves, ostracods, and ﬁsh may be present. The Devonian Orcadian Basin contains wellstudied large open-lake deposits and cycles caused by long-term climatic changes (Donovan 1980).

Fig. 4.8 Characteristics of the Kosi fan, India (from Kelley & Olsen 1993, ﬁg. 23). (Copyright (1993), with permission from Elsevier Science.) (b)

Fig. 4.9 Characteristics of braid plains: (a) section (from Davis, 1992, ﬁg. 3.50) and (b) oblique view (Shyok river at Khapalu, Baltistan Himalaya, Pakistan).

Major environmental complexes and their recognition 77

Fig. 4.10 Meandering stream characteristics (from Tucker 1991, ﬁg. 2.83). (Reproduced with permission of Blackwell Publishing Ltd.)

Fig. 4.11 Characteristics of incised (anastomosing) streams (from Collinson 1986, ﬁg. 3.29). (Reproduced with permission of Blackwell Publishing Ltd.)

78 Chapter 4

Fig. 4.12 Characteristics of lakes (from Talbot & Allen 1996, ﬁg. 4.5a).(Reproduced with permission of Blackwell Publishing Ltd.)

Fig. 4.13 Nearshore carbonate sediments of temperate lakes (from Talbot & Allen 1996, ﬁg. 4.10). (Reproduced with permission of Blackwell Publishing Ltd.)

4.3.3 Wind-dominated systems Wind-dominated systems are fed by winds and ﬂash ﬂoods, with sediments laid down by combinations of mass ﬂows, water, and wind. Periodic debris ﬂows and intermittent water ﬂows deposit immature debris-ﬂow and streamﬂood deposits on alluvial fans which pass basinward into sand dunes and playas (Fig. 4.14). Sediment is reworked by the wind due to restriction of water ﬂow and vegetation to speciﬁc locations or times, in either hot or cold dry environments (Kocurek 1996). Chemical deposits form by evaporation in temporary lakes (which may be large, for example Lake Eyre in Australia) and in soils. Biological systems are unimportant. In a typical wind-dominated system, alluvial fans grade basinward into ephemeral stream and tem-

porary lake deposits and in larger systems into extensive wind-deposited sand seas (ERGS) or deﬂation pebble lags (SERIR). ALLUVIAL FANS in dry climates are steeper and coarser than those in wet regions. The fans are dominated by coarse mass-ﬂow and streamﬂood gravels near their apex and grade downfan into ﬁner grained, braided, and sheetﬂood gravels and sands (Fig. 4.15). EPHEMERAL STREAMS form when ﬂoods extend across the desert ﬂoor from fan channels. The deposits consist of lensitic extensive parallel laminated sheet sands and ﬁne sandy breccias, sometimes overlain by temporary lake silts and clays which may have thin eolian sand laminae formed by migration of dunes across playas, leaving coarse sand lags. Saline lakes and PLAYAS form by evaporation of water

Major environmental complexes and their recognition 79

Fig. 4.14 Intermontane desert environment in Death Valley, USA (from Brookﬁeld 1989, ﬁg. 24).

Fig. 4.15 Idealized cross-section through an arid alluvial fan (from Rust & Koster 1984, ﬁg. 3).

from areas supplied by surface or subsurface water. There are many evaporative lake basins in dry climates, ranging from those in desert rifts and intermontane valleys (e.g. Lake Natron in Tanzania and the Great Salt

Lake in Utah), to those on cold dry cratons (e.g. those in the Dry Valleys of Antarctica). Playas are normally ﬂat dried-out lake beds where laminated and sometimes graded clastics alternate with chemical precipitates and

80 Chapter 4

(b)

(c)

Fig. 4.16 Saline lake deposits: (a) sediment types (from Talbot & Allen 1996, ﬁg. 4.32) (Reproduced with permission of Blackwell Publishing Ltd.); (b, c) Lake Eyre, Australia: (b) coastline and (c) gypsum polygons.

eolian sands. In small rift and intermontane basins, saline lake and playa deposits grade laterally into ephemeral stream and alluvial fan deposits; while in larger, ﬂatter desert areas they grade into (and are interbedded with) eolian deposits (Fig. 4.16).

Sand seas form in dry structural depressions. They comprise sand dunes alternating with interdune ﬂats and are often ﬂanked by arid alluvial fan and ephemeral stream deposits. Large areas of residual lag gravels (serir) occur in places. The largest sand seas occur at the

Major environmental complexes and their recognition 81

Fig. 4.17 Dune types in El Gran Desierto, Mexico (from Brookﬁeld 1992, ﬁg. 1). (Reproduced with permission of the Geological Association of Canada.)

Tropics of Cancer and Capricorn, and in intermontane basins in Central Asia. Depending on wind regime and sand supply, different dune types can occur in different deserts and in different parts of the same desert (Fig. 4.17). Very regular dune patterns may be related to dynamic atmospheric structures. Dunes usually consist of cross-bedded, very wellsorted, well-rounded quartz sands. The type and complexity of the cross-bedding may give some indication of the size and type of dune that formed it. These crossbedded units may be interbedded with interdune sediments of various types, forming distinctive associations (Fig. 4.18). Because the atmospheric boundary layer is so thick (up to several kilometers), eolian dunes may be orders of magnitude larger than water-laid dunes, except in the deeper parts of continental shelves, where submarine bars may approach eolian sizes. LOESS is wind-blown silt. It forms thick, massive and laminated silt deposits with intermittent paleosols in North America and Eurasia, south of the Quaternary glacial limits (Fig. 4.19a). Most seems to come from immature Quaternary glacial deposits, and was blown south to form thick deposits by glacial katabatic winds. Loess deposits have not often been recognized in ancient sedimentary rocks, but they should exist, particularly during glacial epochs. All these continental environments can pass into one

another. For example, in Central Asia, mountain glaciers feed large rivers that extend far into very arid deserts (Fig. 4.19b).

4.4 Environments under water Environments under water differ from ﬂuvial environments on land: ﬁrst, because transport and deposition is mostly by movements within a miscible ﬂuid (in this respect, like wind); and second, because water almost permanently bathes the grains. Marine sediments resemble lake sediments, apart from the differing chemistries and extent. Four main divisions include most environments under water: shorelines, continental shelf, continental slope, and deep sea. In the past, shallow seas existed in the interior of continents (epicontinental or epeiric seas) but are present now only as relatively small, almost-enclosed bodies in temperate and polar areas (e.g. Hudson Bay, Caspian Sea). Clastic, carbonate, or chemical deposits dominate different areas depending on sediment supply and climate. Biological input as sediment is usually unimportant except for carbonates, where the taxonomy of the biotas determines the basic mineralogy and grain size of the sediment.

82 Chapter 4

Fig. 4.18 Interpretation of sections of desert sediments (from Brookﬁeld 1992, ﬁg. 15). (Reproduced with permission of the Geological Association of Canada.) (a)

(b)

Fig. 4.19 (a) Loess section near Lanzhou, northern China; (b) Mt Kongur (7.7 km high) from Tashkurgan valley to south – mass ﬂow, glacial, river, lake, and wind deposits.

Major environmental complexes and their recognition 83

Fig. 4.20 Ternary classiﬁcation of coastal systems (from Reading & Collinson 1996, ﬁg. 6.17). (Reproduced with permission of Blackwell Publishing Ltd.)

4.4.1 Clastic systems Clastic systems are dominated by land-derived sediment. Mixed traction and suspension coastal and shelf areas pass, via mixed slump, turbidity current, and suspension depositional slopes, into suspension-dominated deeper water areas. Clastic shorelines are controlled by the relative importance of sediment supply, tides, and wind-driven currents, and can thus be divided into three extreme end-members: river-dominated, tide-dominated, and wave-dominated environments (including storm surges) (Fig. 4.20). Organisms in these systems provide little sediment. The high productivity of clastic shorelines may allow large numbers of organisms to live, but their remains are greatly diluted by the clastic sediment, and only occasionally, when the sediment supply is cut off for some time, or when already-deposited ﬁne sediment is winnowed, can shell beds accumulate. 1 RIVER-DOMINATED COASTLINES contain birdfoot-type deltas where distributaries build out unconstrained by tide and wave removal of sediments, forming delta-

platform, delta-slope, and prodelta deposits. Deltaplatform deposits consist of linear cross-bedded channel sands, separated by shales, silts, and ﬁne-graded and rippled sands of freshwater to brackish interdistributary bays. These pass seaward into delta-slope silts and then into marine prodelta clays. Although any river-dominated delta section should generally coarsen upwards, the haphazard build-out and crevassing of channels gives a certain randomness to the sediment arrangements, as does the ubiquitous slumping at the delta front (Fig. 4.21). 2 Tide-dominated coastlines contain estuaries with linear sand bodies passing seaward into shore-parallel sand and muds without a marked break in slope. If net deposition is greater than erosion, then a prograding estuary (tide-dominated delta) forms, in which overlapping cross-bedded sand-bar deposits gradually overlie prodelta silts and clays. These, in turn, may be overlain by prograding tidal-ﬂat deposits (Fig. 4.22). 3 Wave-dominated coastlines occur where waves rework river-supplied sediment, forming classic Niletype deltas, and open or barrier coasts (depending on the slope). Nile-type deltas consist of beach and barrier

84 Chapter 4 (a)

Fig. 4.21 Birdfoot deltas: (a) the Mississippi delta from space; (b) distributary channel environments (from Anstey & Chase 1979, ﬁg. 14.2); (c) a vertical section through a birdfoot delta (from Coleman & Prior 1980, ﬁg. 55).

fronts separating prodelta and delta plain areas (Fig. 4.23). Barrier coasts have beaches and barriers with shoreline-parallel belts of reworked sands cut by tidal channels, passing shoreward into spillover sands, lagoonal silts, clays, and marshes, and passing

seaward into shallow marine cross-bedded and graded sands (often storm-deposited), silts, and clays (Fig. 4.24). Barrier complexes form on abandoned deltas as the subsiding delta platform is reworked by waves (Fig. 4.25).

Major environmental complexes and their recognition 85 (a)

(c)

(b)

Fig. 4.22 Tide-dominated deltas: (a) mouth of the Irrawaddy, Burma (courtesy NASA); (b) the Fly delta, New Guinea (from Anstey & Chase 1979, ﬁg. 14.1C); (c) a vertical section through a tide-dominated delta (from Coleman & Prior 1980, ﬁg. 57).

In prograding deltas, as in most other environments, the boundaries between the different sedimentary units are not time-lines. At any one time, nearshore sediments pass seaward into deeper water sediments. The depositional surfaces are also usually at small angles to the

horizontal, except in some coarse-grained systems such as arid alluvial fans (Fig. 4.26). CONTINENTAL SHELVES slope gently from the shoreline to depths of about 100–250 m, which normally marks the shelf/slope break. They range in width from a few to

86 Chapter 4 (a)

(c)

(b)

Fig. 4.23 Wave-dominated deltas: (a) the Niger delta from space (courtesy NASA); (b) the Niger delta (from Anstey & Chase 1979, ﬁg. 14.1); (c) vertical section through a wave-dominated delta (from Coleman & Prior 1980, ﬁg. 61).

several hundred kilometers and are characterized by free connection with the open ocean (in contrast to epicontinental seas). The dominant processes are tidaland wind-driven currents, including storm waves and

storm-induced underﬂows, but glaciers can extend across the shelves at times. Tide-dominated shelves consist of various types and sizes of cross-bedded sand bodies (depending on the

Major environmental complexes and their recognition 87 (c)

(a)

(b) (d)

Fig. 4.24 Wave-dominated barrier coasts: (a) offshore barrier, Mississippi delta, Chandeleur Islands (courtesy NOAA); (b) transgressive barrier complex; (c) transgressive barrier section; (d) barrier-inlet section. (b–d from Reinson, 1992, ﬁgs. 3, 9.) (Reproduced with permission of the Geological Association of Canada.)

tides), which pass laterally into interbedded cross-bedded sands, silts, and clays (Dalrymple 1992). Some deeper, tide-dominated shelf sand bodies resemble desert dunes: both are formed by deep, wide ﬂuid ﬂows. Wave- and storm-dominated shelves consist of hummocky cross-bedded and graded and rippled sands,

which are thin, ﬁne, and pass seawards to be interbedded with silts and clays (Fig. 4.27). Glacially inﬂuenced shelves, like those around Antarctica, are characterized by dropstones, submarine moraines, and outwash fans (Fig. 4.28). Finer clastic and biochemical sediments, deposited during warmer

88 Chapter 4

(b)

(c)

Fig. 4.25 (a) Barrier complex of an abandoned Mississippi delta lobe (from Elliot 1986, ﬁg. 6.27); Chandeleur Island (b) before and (c) after Hurricane George in 1998 (courtesy USGS).

interglacial periods, may be interbedded. Rapid glacially controlled changes of relative sea level greatly affect such successions (see Chapter 11). All shelves tend to converge to an equilibrium proﬁle when sediment supplied at the shoreline is transported across the shelf into the deeper water continental slope.

Sedimentation should thus be spasmodic and slow after equilibrium, and controlled by the slow rates of subsidence characteristic of mature shelves and by relative sea-level changes. Modern shelves, however, are poor analogies for ancient shelves since the former are in disequilibrium after the rapid postglacial rise in sea level of

Major environmental complexes and their recognition 89

Fig. 4.26 Time-lines and sediment type in a prograding delta (great vertical exaggeration) (from Coleman & Prior 1980, ﬁg. 26).

(a)

(b)

Fig. 4.27 Wave- and storm-dominated shelves: (a) processes; (b) vertical section (from Walker & Plint 1992, ﬁgs 14, 21).

the last 10,000 years. The resulting palimpsest distribution of relic sediments cannot be directly related to modern processes (Fig. 4.29). CONTINENTAL SLOPES are transitional between shelf and deep-sea environments and coincide with the structural change from continental to oceanic crust. They are similarly transition zones between wave- and tidedominated shelves and the mass-ﬂow and currentdominated deep oceans (Fig. 4.30). Huge circulation systems often run along slopes, reworking sediment on a grand scale. Modern slope successions are poorly known because few cores have been recovered. Most

information comes from seismic records and inferred processes. Apart from deposits formed by simple settling of suspension, slopes seem to be characterized by mass-ﬂow deposits (often of enormous size) and “contourites”. Mass ﬂows involve either whole segments of the slope or consist of smaller ﬂows channelized along canyons incised in the slope. CONTOURITES are usually ﬁne sediments deposited from slope-hugging oceanic currents, but can be coarse cross-bedded sands. Where sedimentation is slow, organisms can form an important part of the sediment. Shelf-edge patch accumulations of epifauna are common, and are

90 Chapter 4

Fig. 4.28 Glacially inﬂuenced shelves: (a) processes (after McCabe et al. 1984); (b) vertical section during glacial advance and retreat (from Bennet & Glasser 1996, ﬁg. 10.11).

sometimes large enough to be labelled “reefs” (Fig. 4.31). Deep-sea deposits cover over 60% of the earth, but are little known. Most sediment accumulates at the continental-slope/deep-sea transition as deep-sea fans, which grade out into deep-sea plain condensed pelagic sediments.

1 DEEP-SEA FANS form when turbid suspensions slide down continental slopes, and decelerate on reaching the deep-sea ﬂoor. Here they form submarine fans much larger than the fans found on land. The sediments resemble those of subaerial fans, with similar channel and overbank systems, except that the ﬂow can mix with the ambient ﬂuid (water) (Fig. 4.32).

Major environmental complexes and their recognition 91

Fig. 4.29 Relic sediment on Recent Alaska continental shelf (from Carlson et al. 1990, ﬁg. 6A). (Reproduced with permission of the Geological Society.)

Fig. 4.30 Hydrodynamic regimes in the oceans (from Einsele 1992, ﬁg. 1.5). (Copyright (1992). Reproduced with permission of Springer-Verlag.)

The largest fans are those found in front of the deltas of large rivers such as the Amazon and the Indus, which consist of mature sediments. Smaller fans draining active orogens such as the Andes consist of less mature sediments and occur in front of smaller rivers.

2 Deep-sea (abyssal) plains lie between continental slopes and oceanic ridge and arc systems, and may be thousands of kilometers long and hundreds of kilometers wide (Fig. 4.33a). Sediments are generally thin and composed of the ﬁnest and furthest-transported clastic sediment, as well as true pelagic sediments. Sections of

(a)

(b)

Fig. 4.31 Deep, cold-water, shelf-edge Lophelia reefs: (a) ridge and reef off Trondheim, Norway (from Freiwald & Wilson 1998, ﬁg. 2); (b) Lophelia reef (courtesy of E. A. Moen). (a)

Fig. 4.32 Submarine fans: (a) view (from Reineck & Singh 1973, ﬁg. 541) (Copyright (1973). Reproduced with permission of Springer-Verlag.); (b) sub-environments; (c) vertical section of a prograding fan (b, c from Walker 1992, ﬁgs 15, 20).

Major environmental complexes and their recognition 93 (a)

(b)

Fig. 4.33 Deep-sea plains: (a) Sohm and Hatteras abyssal plains, western Atlantic (from Pickering et al. 1989, ﬁg. 8.6)(Reproduced with permission of HarperCollins.); (b) ancient deep-sea plain deposits (from Stow 1986, ﬁg. 12.34). (Reproduced with permission of Blackwell Publishing Ltd.)

deep-sea plain deposits consist of varying proportions of ﬁne-grained distal turbidites, volcanic ashes, mudstones, radiolarian cherts, and foraminiferal and other chalks and limestones (Fig. 4.33b). 4.4.2 Carbonate systems Carbonate systems develop where the supply of landderived clastics is reduced, and biologically and chemically produced carbonate can accumulate without dilution. The critical relationship is the rate of supply of clastics versus the rate of production (or supply) or

carbonates. There are, of course, all transitions between clastic- and carbonate-dominated systems, though the different sediments tend to alternate rather than be present in equal quantities. Modern carbonate systems occur in all tectonic situations, from poles to tropics. Warm, shallow-water environments should not be assumed for any ancient carbonates without proof (Fig. 4.34). There is also a tendency to equate clastics with wet hinterland conditions and carbonates with dry, desert hinterland conditions; but this is not necessarily the case. The Sunda shelf of Southeast Asia is now a warm tropical shelf with a high clastic sediment input. In the

94 Chapter 4

Fig. 4.34 Modern carbonate sediments, distribution, and character (from Copper 1992, ﬁg. 4). (Reproduced with permission of the Geological Association of Canada.)

past, a lower clastic input due to higher sea levels allowed vast areas to develop carbonate sediments. Many carbonate systems resemble clastic systems, except that the input of sediment is from organisms and solutions within the environment. Since most carbonate sediments form subaqueously (clastic limestone, detrital calcrete soils, and semi-arid precipitates are exceptions), most carbonate sediments, both in seas and lakes, accumulate in shoreline to deep-basin situations. Carbonate shorelines, like clastic shorelines, are controlled by the relative importance of sediment supply, tides, and winds; and consist of tidal ﬂats, lagoons, barrier bars, and open coastlines. Tidal ﬂats have supratidal sabkhas, which undergo early dolomitization, and algal mats are important along the shorelines. Gypsum laminae alternate in places with peloids formed from the fecal pellets of ostracods and brine shrimps. Lagoons tend to have high and variable salinities and are dominated by low-diversity, but numerically abundant,

faunas of tolerant forms such as ostracods and molluscs. Barriers can comprise bioclastic sands, comparable to clastic barriers, or coral reefs. Barrier inlets in warm water often have oolitic bars (Fig. 4.35). Carbonate shelves occur in tropical to temperate areas and can be divided into ramps and rimmed shelves. RAMPS show a gradual deepening towards the continental slope and are swept by waves because there is no marginal barrier. Waves tend to break at speciﬁc depths offshore. So there may be a wave-dominated bioclastic barrier separating a protected area from the deeper ramp. The climate determines the biotas that form the carbonate and some of the details of sedimentation (Fig. 4.36). RIMMED SHELVES have coral reefs, sand shoals, or islands along their seaward margins, which effectively reduce connection with the open ocean. The sediment types are dependent on the shelf depth, climate, and the type of organisms. If carbonate sedimentation keeps pace with

Major environmental complexes and their recognition 95 (a)

(b)

Fig. 4.35 Carbonate shoreline environments: (a) shuttle view of sabkha, lagoons, and inlets of the southern Persian Gulf (courtesy NASA); (b) shallow carbonate banks, Bermuda; (c) middle sabkha section; (d) shoreline algal mat section; (e) channel section (c-e from Pratt et al. 1992, ﬁg. 18) (Reproduced with permission of the Geological Association of Canada.)

subsidence or sea-level rise, shallow sands and muds pass offshore into sands, reefs, and bars. If not, then deep-water shelves may be ﬂoored with muds (Fig. 4.37).

REEFS are notoriously difﬁcult to deﬁne because they mean different things to different people. To mariners, the term means any rock structures at or just below the sea surface that are a hazard to navigation. To biologists,

96 Chapter 4

Fig. 4.36 Warm- and cold-water ramps (from Jones & Desrocher 1992, ﬁg. 15) (Reproduced with permission of the Geological Association of Canada.)

the term means rich biological communities that form rigid wave-resistant structures in shallow tropical seas. To geologists, the term means a massive three-dimensional carbonate structure that is surrounded by bedded carbonates and that may have distinctive fossils. The big problem is wave-resistance and temperature. Large Recent “reefs” also occur at depth and in cold water. Furthermore, wave resistance can only rarely be demonstrated in fossil masses, and contemporary depth and temperature conditions are usually very difﬁcult to infer. The neutral term “bioherm” can be used for any lens-shaped carbonate mound. Large barrier reefs may build out seaward over deeper water sediments, and may

isolate marine basins sufﬁciently that a drop in sea level will allow an evaporitic basin to form (Fig. 4.38). CARBONATE SLOPES may be much steeper than clastic slopes because of the rapid growth of reefs and early cementation. Mass ﬂows down reef fronts form thick breccias, graded gravels and sands, and thin turbidite sands interstratiﬁed with suspension muds (Fig. 4.39). Deep-sea carbonates are formed by the accumulation of shells of planktonic organisms. Recent calcareous plankton have high production rates in warm, low-latitude surface waters, mostly within 60° of the equator. Siliceous plankton are commoner in cooler surface waters, especially in nutrient-rich areas such as

Major environmental complexes and their recognition 97

Fig. 4.37 (a) Shallow, warm-water rimmed shelf; (b) deep, warm-water rimmed shelf (from Jones & Desrocher 1992, ﬁg. 20). (Reproduced with permission of the Geological Association of Canada.)

upwellings. Only siliceous oozes accumulate below the carbonate compensation depth, which is currently several kilometers, but which has varied in the past. Deep-sea carbonates accumulate at about 0.01–0.05 mm/year, which is low compared to shelf carbonates. Since all deep-sea deposits eventually end up crushed into subduction zones, ancient deep-sea carbonates are almost entirely conﬁned to orogenic belts. 4.4.3 Epicontinental (epeiric) seas EPICONTINENTAL (EPEIRIC) SEAS are extensive shallow seas that covered vast areas of continents at particular times in earth history. The Middle Ordovician seas of North America stretched, almost unbroken by land, across the entire continent. Such seas are difﬁcult to interpret

because no modern counterparts exist. Only the Baltic Sea and Hudson’s Bay are comparable; yet these are much smaller and in cold seasonal climates. Ancient epicontinental seas were often large and extensive, had low average bottom slopes, and were frequently partially isolated from ocean basins. Some major problems are unresolved. Were tides important or negligible? Did storms move most of the sediment around? Were salinities mixed or strongly controlled by local river input or evaporation? Were organisms limited by nutrient turnover? Was the growth of shelled organisms limited by carbonate recycling? Studies using modelling and/or observations of ancient epicontinental seas are often ambiguous or contradictory. For example, in the Devonian Catskill Sea of eastern North America, numerical modelling showed

(a)

Fig. 4.38 Permian El Capitan reef of USA: (a) view from basin; (b) map and cross-section. (Reproduced by kind permission of Mark Eberle.)

Major environmental complexes and their recognition 99

Fig. 4.39 Carbonate slopes and an ideal section of a prograding margin (from Coniglio & Dix 1992, ﬁgs 17, 23). (Reproduced with permission of the Geological Association of Canada.)

that a moderate tidal range was possible in this sea, and some observations are compatible with this (Slingerland 1986). However, the same observational evidence is used to infer deposition by tropical storms without signiﬁcant tides (Craft & Bridge 1987). Perhaps the only sensible course is to study the processes in existing shallow inland seas, such as the Caspian Sea, the Baltic Sea, and Hudson’s Bay, and make what limited extrapolations are possible. In these seas, the fairweather wave base is quite shallow, tides are negligible, and the main process moving sediment is storm waves. The storms can raise or lower the sea level by several meters locally, forming currents that reach 1 m/s (3.6 km/h).

4.5 Mixed environments All the above environments can and do pass laterally into each other and succeed each other in time as conditions change. For example, distinctive limestone often caps glacial tillites in Neoproterozoic successions, and Recent island-arc volcaniclastics may pass very rapidly into reef carbonates, as in the small islands of the Caribbean.

4.6 Peculiar environments Many ancient sediments have no direct modern analogies. Some physical processes are not observable today

100 Chapter 4 or are very rare; for example, catastrophic volcanic eruptions and large meteorite impacts. Biological evolution has altered the composition of shells of organisms; for example, Paleozoic fossils are mostly calcite. In the past, some chemical and physical processes may have

taken place under different conditions than at present, or produced strange lithologies in unknown ways, especially without the inﬂuence of the present biosphere; for example the formation of BANDED IRONSTONE and oolitic ironstone (Mel’nyuk 1982; Young & Taylor 1989).

II Tracing environments in space and time

Hutton’s famous angular unconformity at Siccar Point, Scotland.

After learning how to identify and classify sediments, how to recognize the processes that form individual layers, and how to recognize environments, these capabilities must be applied to actual rocks on and below the land surface. The basic data of stratigraphy (and most other studies in geology) come from studying individual sections and tracing units deﬁned in these sections across a landscape. After the preliminary overview, good section descriptions and the construction and interpretation of maps are the basis for all other studies. However, both section description and mapping are highly interpreta-

tive. So, it is only by describing and mapping rocks yourself that you can appreciate the amount of extrapolation and interpretation (and yes – guesswork!) needed to construct even the simplest geologic map. People with little ﬁeld experience often have an uncritical and unwarranted conﬁdence in published section descriptions, unit classiﬁcations, and maps, and do not realize their limitations. As Blatt et al. (1991) noted, “Many a geologist has visited an area with geologic map in hand, expecting to see a particular stratigraphic relationship or facies change, only to end up in a ploughed wheat ﬁeld or 100 meters from shore in a rowboat.”

5 The vertical dimension

5.1 5.2 5.3 5.4

The local section Breaks in the record Dividing the local section: the type section Strata and stratiﬁcation

Stratigraphers, after a seismic, aerial, or pedestrian overview of an area, start by looking at individual outcrops or boreholes, and describing local sections and the breaks within them.

5.1 The local section Despite the vagueness, incoherence, and downright stupidity of many stratigraphic classiﬁcations, some sort of subdivision of a rock column is essential. Should this be based on rock type or on time? Since rock units, such as sandstones deposited in a shallow sea, can cut across time as sea levels rise and fall, space and time units must be separated (see facies on p. 129). Space- or rockstratigraphic (lithostratigraphic) units are therefore kept distinct from relative time units, which are usually based on fossils or magnetic reversals. Likewise, relative time units are not the same as absolute ages, which are based on radiometric dates. Although some people consider relative and absolute time units as one, they require mutual calibrations which are always changing (e.g.

ages of the base of the Cambrian have ranged from 600 to 535 Ma). For these reasons, three stratigraphic classiﬁcations are normally used: (i) space types (LITHOSTRATIGRAPHIC), which simply summarize the geometry of rocks; (ii) RELATIVE TIME types (CHRONOSTRATIGRAPHIC), which give time equivalence or older/younger relationships; and (iii) NUMERICAL TIME types (GEOCHRONOLOGIC), which give ages in years (Table 5.1). Note that BIOSTRATIGRAPHY is included with chronostratigraphy here. When dividing a local section, make sure you keep these three types of classiﬁcation distinct. Failure to do so has been the cause of much confusion, particularly before the mid-20th century, since these three divisions did not spring up overnight. Combined rock and relative-time (biostratigraphic) subdivisions were used by the ﬁrst classiﬁers of the geologic column: rock and fossil changes were thought to occur at the same time. For example, Werner’s division into successive layers in a “Universal Ocean” was essentially rock-stratigraphic, but the successive layers were assumed to correspond to the time units based on Cuvier’s succession of catastrophic creations and

104 Chapter 5 Table 5.1 Summary of categories and unit terms in stratigraphic classiﬁcation (after Hedberg 1976). Stratigraphic categories

Principal stratigraphic unit terms

Lithostratigraphic Deﬁnes a body of rock strata uniﬁed by overall homogeneity of lithology or combination of lithologies; may be sedimentary, metamorphic, or igneous

Group Formation Member Bed(s)

Biostratigraphic Deﬁnes a body of rock uniﬁed by its fossil content

Biozones Assemblage zones Range zones (various kinds) Acme zones Interval zones Other kinds of biozones

Chronostratigraphic Deﬁnes a body of rock uniﬁed by being formed all rocks during a speciﬁc interval of geologic time; represents formed anywhere during a certain segment of earth history

(Equivalent units)

Geochronologic Deﬁnes a unit of geologic time determined by geologic methods; may correspond to the time-span of a stratigraphic unit

Fig. 5.1 Sedgwick, Murchison, and their ﬁeld areas (see Fig. 5.2).

Eonothem Erathem System Series Stage Chronozone

Eon Era Period Epoch Age Chron

extinctions of faunas. However, the best example is the division of the Lower Paleozoic into the Cambrian and Silurian systems by Sedgwick and Murchison (Secord 1986). This division and the subsequent controversy have always occupied a central place in the history of geology as one of the great set-pieces of the science’s “golden age”, and it is still of interest in emphasizing the ongoing problems of lithostratigraphic (space) versus biostratigraphic (relative time) classiﬁcations of sections. In 1831, two close friends, Sedgwick and Murchison, entered Wales from opposite ends to try to unravel the rock sequences (Fig. 5.1). Sedgwick emphasized structure, the tracing of rock units to show deformations, based on Beaumont’s catastrophic ideas that sudden upheavals of mountain chains were represented in the rock record by time breaks (unconformities) and massive extinctions. His ability to visualize rock masses in three dimensions was perfectly adapted to interpret the complicated, virtually unfossiliferous rocks of North Wales. Murchison, on the other hand, emphasized mapping, the long-distance correlation of rock masses with

The vertical dimension 105 (a)

(b)

(c)

(d)

Fig. 5.2 The landscapes of Wales and the borders (locations shown in Fig. 5.1). (a) The Cambrian mountains: ﬁeldwork as Sedgwick knew it. (b) Snowdon, the disputed volcanics and slates between Murchison’s and Sedgwick’s systems. (c) Wenlock edge, Silurian reef limestone escarpment. (d) Carreg Castle, Llandelio, on tilted Carboniferous limestone.

fossils: fossils that did not change as much laterally as the enclosing rocks. His methods were better suited to the relatively undeformed, poorly exposed, fossiliferous rocks of the Welsh–English border regions. The landscapes of these areas have hardly changed since Sedgwick and Murchison studied them, and you can still visit these areas and grapple with the same problems they had (Fig. 5.2). Sedgwick and Murchison quarrelled about the overlap of their respective systems. Sedgwick’s neglect of fossils restricted the range of his Cambrian system, and Murchison’s Silurian system, based on fossils, reigned supreme. Not until Lapworth’s 1879 insertion of the Ordovician system was the controversy solved; and this solution was not formally accepted until

1960 (Bassett 1979). Both Sedgwick and Murchison realized that their sections were incomplete. Like Hutton, they realized that breaks in the sections recorded long periods of erosion separating periods of deposition (Fig. 5.3).

5.2 Breaks in the record Breaks in the stratigraphic record, or UNCONFORMITIES, were used by Hutton (1795) to conﬁrm his cyclical view of earth history. Apparently less signiﬁcant time breaks, or DIASTEMS, mark pauses in sedimentation or local erosion. Both are caused by relative changes in the base level of erosion.

106 Chapter 5

Fig. 5.3 Cross-section across Wales, with unconformities (heavy wavy lines) separating systems (from Petersen & Rigby 1990, ﬁg. 3.2). (Copyright 1990. Reproduced with permission of the McGraw-Hill Companies.)

The type of unconformity has no time signiﬁcance. Some nonconformities may represent less than one million years, while some paraconformities may represent hundreds of millions of years. For example, in Estonia, soft Quaternary clays rest paraconformably on soft Cambrian clays, and over 400 million years unrepresented between them. The type and time value of unconformities can also change from place to place due to differential uplift and erosion (Fig. 5.6). 5.2.2 Diastems

Fig. 5.4 Quinag, northwestern Scotland: Late Precambrian sandstones unconformable on a high-grade metamorphic basement.

5.2.1 Unconformities Unconformities involve relative uplift and erosion followed by renewed deposition, often under different environmental conditions (Fig. 5.4). Four different types of unconformity are often distinguished (Fig. 5.5): (i) a nonconformity has sedimentary rock on an igneous and/or metamorphic basement; (ii) an angular unconformity has sedimentary rock on tilted, truncated, layered rocks; (iii) a disconformity has sedimentary rock on a rolling eroded surface; and (iv) a paraconformity has parallel layers above and below the unconformity. A paraconformity can often only be detected by dating the rocks above and below the unconformity, and noting the time gap.

Diastems are supposedly more minor time breaks due to variations in the normal processes in an environment, and these are of several scales and types. Before damming, the River Nile in Egypt ﬂooded for several weeks each year, depositing a thin silt layer on its ﬂoodplain (Fig. 5.7a). For the rest of the year the silt was exposed, subject to weathering and colonization by plants and animals. So, the actual silt layers in the ﬂoodplain section represent only a few weeks out of every year, and between each layer is a NON–DEPOSITIONAL DIASTEM representing the rest of the year. Some events during the diastem may be preserved as plant rootlets, animal traces, and so on, but most are lost. On a larger scale, silt sections can be partly or even entirely removed by the migration of river channels, forming an even longer EROSIONAL DIASTEM (Fig. 5.7b). Analogous diastems can be seen in ancient deltas (Fig. 5.8). On a larger scale, entire deltas can shift, building up overlapping lenses separated by even longer diastems (or do we call these minor unconformities?) (Fig. 5.9). Non-depositional and erosional diastems are caused by changes in the BASE LEVEL OF EROSION OR AGGRADATION,

The vertical dimension 107

Fig. 5.5 Types of unconformity (from Dunbar & Rodgers 1957, ﬁg. 57). (Author given permission.)

Fig. 5.6 Variation in unconformity type at the base of the Pennsylvanian in Wyoming (from Dunbar & Rodgers 1957, ﬁg. 62) (Author given permission.) (a)

(b)

Fig. 5.7 Nile diastems: (a) non-depositional diastem, ﬂoodplain; (b) erosional diastem, migrating channel.

108 Chapter 5

Fig. 5.8 Erosional (channel) and depositional (soil) diastems in a Carboniferous coal-bearing succession, Kentucky, USA (modiﬁed from Horne et al. 1978, ﬁg. 9).

Fig. 5.9 Successive overlapping delta lobes of the Mississippi delta. Numbers 1–7 are the locations of the successive lobes (from Scruton 1960, ﬁg. 15).

The vertical dimension 109

Fig. 5.10 Base level of erosion and aggradation in a shallow shelf environment (from Dunbar & Rodgers 1957, ﬁg. 68). (Author given permission.)

Fig. 5.11 Diagram for diastems and unconformities (modiﬁed from Barrell 1917, ﬁg. 5). (Copyright © 1917. Reproduced with permission of the publisher, the Geological Society of America, Boulder. Colorado, USA.)

due to ﬂuctuations in the environment; for example on a shallow shelf (Fig. 5.10). During storms, nearshore erosional diastems develop, while the sediment transported offshore accumulates as storm layers. As the storm wanes, sedimentation resumes in the nearshore environment and declines or ceases in the offshore environment where non-depositional diastems may form. Any variable or intermittent process gives such diastems. Diastems and their effects were most clearly explained by Barrell (1917) using the concept of changing base

levels of erosion (Fig. 5.11). In Fig. 5.11, Curve A–A represents the overall base-level change due to long-term changes in sea level, rate of subsidence, and rate of sedimentation. Curve B–B represents the rise and fall of base level due to shorter-term changes; for example, sea-level changes due to glaciation and deglaciation. Curve C–C represents the ﬁnal pattern of movement of base level when still shorter-term ﬂuctuations are superimposed; for example, periodic major storms. Note that C–C is still highly simpliﬁed from the normal situation. Above the

110 Chapter 5

Fig. 5.12 One of Brinkman’s (1929) diastem time values (from Dunbar & Rodgers 1957, ﬁg. 69). (Author given permission.)

diagram, the black bars show permanently preserved sediment; dotted bars show sediment temporarily preserved during base-level rises, but which is removed by major base-level falls. Erosional diastems occur when ‘C’ base levels drop. On the left, the total amount of sediment deposited and ﬁnally preserved is shown for the three base-level cycles X, Y, and Z. Note that less than half of the total in each is usually preserved. Curves A, B, and C pick out major unconformities, minor unconformities, and erosional diastems, respectively; however, there is no conceptual difference among them. The time value of individual diastems is almost impossible to work out. Brinkman (1929) tried it with

Oxfordian ammonites (Jurassic) from an English brickpit (Fig. 5.12). He noted that characteristics such as rib number and bifurcation of ribs gradually changed up the section; apart from sudden jumps at levels with squashed ammonite concentrations (e.g. at A in Fig. 5.12). By pulling back the graphs to show a linear change in characters, Brinkman estimated the relative amount of time missing at various levels (A’). By working out the total amount of time represented by the brickpit succession, he could calculate the actual time value of the diastems. Unfortunately, not only is the dating of Jurassic stages particularly poor, but the linear evolution assumed by Brinkman may not be correct

The vertical dimension 111

Fig. 5.13 Interglacial terraces and glacial valley cutting, lower Mississippi River, USA (from Strahler 1971, ﬁg. 41.18).

The canyon is now ﬁlled with sediments that were deposited as base levels rose with rising Mediterranean sea levels (Fig. 5.15). However, the canyon, together with other evidence, indicates that the Mediterranean was dry around 6 Ma.

5.3 Dividing the local section: the type section

Fig. 5.14 River terraces on the Toze River, Zanskar Himalaya, India.

(Raup & Crick 1982). Note also the enormous scatter of the points around the regression lines. Unconformities and diastems can occur at many scales, and may repeat regularly. For example, repetitive Quaternary river terraces along the Mississippi, and their associated sediment, were formerly attributed to regular base-level changes caused by glacio-eustatic ﬂuctuations in sea levels, superimposed on slow general uplift (Fig. 5.13). This interpretation is now less certain, but regular terraces along internally draining rivers in Central Asia are still attributed to interactions between tectonic uplift of the adjacent Tien Shan and Quaternary climatic change (Fig. 5.14). Very unusual, even CATASTROPHIC CHANGES in Mediterranean sea levels were needed to explain the Nile canyon. This canyon is cut 250 m below sea level at Aswan, almost 1000 km from the mouth of the Nile.

Local sections need subdividing into manageable units in order to describe them. To start with, lithological subdivisions are usually made on characteristics observable in the ﬁeld, such as rock type, fossil content, color, and so on. At the same time, suitable fossils may be collected, magnetic directions measured, and so on, for relative and numerical dating (see Chapter 7). In the 19th century, a fairly informal naming system of rock units was used, but this has now been superseded by more formal methods, usually outlined in the Stratigraphic Codes of various countries (Fig. 5.16). However, division of a section into formal stratigraphic units is often difﬁcult, is subject to personal preference, and still remains something of an art. The main purpose of stratigraphic division is to make it easier to understand the strata – an aim often forgotten. The basic lithostratigraphic unit of subdivision normally used today is the FORMATION. A formation is a homogeneous rock unit, or an association of distinct interbedded rock units, which are separable from the rock units above and below, and which can be shown on a geological map of at least 1 : 50,000 scale. A formation may be homogeneous sandstone, or interbedded sandstone, shale and limestone, or any other combination, and can be any thickness. In some areas, it might be useful to have formations only a few meters thick; for

112 Chapter 5

Fig. 5.15 Sediment-ﬁlled canyon below the Nile River at Aswan, Egypt (from Hsu 1972, p. 36).

Fig. 5.16 19th and 20th century Upper Mississippi Valley rock units (redrawn from Krumbein & Sloss 1963, ﬁgs 2.2, 2.3). (© 1951, 1963 by W. C. Krumbein and L.L. Sloss. Used with the permission of W. H. Freeman and Company.)

The vertical dimension 113

Fig. 5.17 Time-transgressive Bridport Sandstone, southern England (from Krumbein & Sloss 1963, ﬁg. 10.12). (© 1951, 1963 by W. C. Krumbein and L.L. Sloss. Used with the permission of W. H. Freeman and Company.)

Fig. 5.18 Member, tongue, and lens of sandstone in a mixed sandstone–shale formation.

Fig. 5.19 Alternating limestone and shale in the classic Lower Lias (Jurassic) of Dorset, England.

114 Chapter 5 example, where rock sections are condensed. Elsewhere, homogeneous rock successions might have formations thousands of meters thick; for example, where sections are made up of uniform thick delta clays. Formation boundaries are called contacts, and are normally placed where marked changes in rock types occur; for example, where limestone changes to clay and sandstone. Contacts can also be placed at an arbitrary marker bed in a gradational succession. There are two conﬂicting concepts of formations: 1 Formations should, as far as possible, have components that are time equivalents. 2 Formations are strictly rock units with no requirement for time equivalence. The second concept is more acceptable today since, in the past, attempts to deﬁne time-equivalent formations caused miscorrelation and confusion. Rock and time units should be kept distinct. For example, the Jurassic Bridport and related sandstones of England form a sheet sandstone unit that is entirely older in the north than in the south (Fig. 5.17). The time-transgressive nature of this formation is the result of a slow rise in sea level. Some of the problems of tracing formations are discussed in the next chapter. Formations are named after a locality where that formation is particularly well exposed, and where its character is best seen. This is the TYPE SECTION. Formation names consist of two parts: the ﬁrst part is the locality name; the second is usually the dominant rock type. For example, the Navajo Sandstone is named after a tribe occupying the area where the rock is best exposed and after the dominant rock type. If there is no dominant rock type, a locality name is followed simply by the word “formation,” for example the Carmel Formation. For discussion of larger areas and units, formations may be assembled into GROUPS, which deﬁne even larger rock complexes. For example, the Navajo Sandstone, an eolian sand deposit, forms part of the Glen Canyon

Group, which contains formations deposited in predominantly arid environments. This group differs from the underlying limestone (shallow sea) and overlying sandstone and shale (river ﬂoodplain) formations. Formations can be subdivided to describe variations within them (Fig. 5.18). A MEMBER is a laterally persistent rock unit within a formation. A TONGUE dies out laterally in one direction. A LENS dies out in both directions.

5.4 Strata and stratiﬁcation The most obvious feature of layered rocks not yet mentioned is that they usually occur in distinct layers or STRATA, i.e. they are usually stratiﬁed. STRATIFICATION is fundamental to stratigraphy, but the cause of much stratiﬁcation is obscure. The majority of bedding planes are probably caused by breaks in sedimentation of different duration. Sand moving over ripples and dunes periodically avalanches down the lee slopes, forming the distinct thin layers of cross-laminae and cross-bedding. Waning ﬂows of turbidity currents form distinct graded beds with sharp bedding planes above and below. Compression of clays forms the layering of shales, which may also show differences in layers due to chemical differences. Variations in carbonate and other chemical precipitates may lead to layers of differing composition. Erosion down to cemented beds may allow boring and colonization of the surfaces separating layers, while variations in intensity and type of burrowing can also form distinct layers. Stratiﬁcation can thus be produced by a variety of physical, chemical, and biological processes during deposition. Diagenetic segregation may produce misleading strata unrelated to the original layering. For example, diagenetic carbonate segregation is at least partly responsible for some interbedded limestone–shale successions (Fig. 5.19).

6 The horizontal dimension

6.1 Physical correlation 6.2 Lateral changes 6.3 Mapping

Eventually, you must trace the stratigraphic features of one local section laterally into other local sections – a procedure known as correlation. There are two fundamental ways of doing correlation: in space or in time. Both space and time correlations can be precise or imprecise (Fig. 6.1). Precise, but inaccurate, methods are sometimes difﬁcult to evaluate: radiometric dating can give apparently exact dates that do not date the events being studied. Precise space correlations can be made with the unique characteristics of an environment. For example, low-angle beach cross-bedding may be traced within a sandstone body, marking the successive positions of the beach. Imprecise space correlations are more usual. For example, an interbedded limestone and shale unit may be traced from one area to another, even though individual beds cannot. Precise time correlations can be made using unique layers deposited during unusual events. Precise correlation of the end of the Cretaceous is possible due to a unique worldwide impact layer. Individual ash layers

from volcanoes can be used if speciﬁc ashfall layers can be identiﬁed. Imprecise time correlations are done with fossils, magnetic reversals, and other more restricted methods. These are imprecise because the fossil and magnetic changes occur over variable time intervals; they are not instantaneous. This chapter covers space (or physical) correlation, lateral changes, and their mapping. Time correlation is covered in Chapter 7.

6.1 Physical correlation Space correlation involves tracing some physical aspect or relationship of the strata; tracing unique rock and/or fossil types, characteristic rock sequences, geochemical features, unconformities, and so on. A few of these methods also give a precise time correlation, for example unique ash layers. The bases for physical correlation are those proposed in the 17th century by Nicolaus Steno (Nils Steensen), a

116 Chapter 6

Fig. 6.1 Target precision, accuracy, and resolution. One cluster has high precision (high reproducibility) but is inaccurate; another is more accurate (closer to the centre) but imprecise (poorly clustered). If the clusters are distinct, resolution is higher than if the clusters overlap (from Prothero & Schwab 1996, ﬁg. 18.16).

Fig. 6.2 The Grand Canyon: superposition, initial horizontality, and original continuity.

Danish physician who traveled widely and studied the rocks of Tuscany in northern Italy. In 1669, Steno proposed three laws (or principles) which form the basis for the physical correlation of rocks: superposition, initial horizontality, and original continuity. The Grand Canyon of the western USA exempliﬁes these principles (Fig. 6.2). The PRINCIPLE OF SUPERPOSITION states that strata become younger upwards – unless the entire section has been overturned during deformation. Although not exactly an earth-shattering conclusion, this has nevertheless been disputed at times, especially by paleontologists, who sometimes thought that layers with unusual fossils were deposited in caves. The PRINCIPLE OF INITIAL HORIZONTALITY states that strata were initially deposited ﬂat. Although granular sediments may pile up until they collapse to an angle of

repose (as in gravel bars and sand dunes), most sediments are deposited as subhorizontal layers. If much tilted then they have been disturbed since deposition. The PRINCIPLE OF ORIGINAL CONTINUITY states that rock layers were originally continuous. For example, horizontal strata exposed in river valleys were once connected and have since been eroded by the rivers cutting them. Steno also used the PRINCIPLE OF CROSS-CUTTING RELATIONSHIPS, which states that a tabular layer is younger than any layers it cuts across. Around Florence, Italy, Steno also saw inclined marine sedimentary rock cut across and overlapped by softer horizontal marine sediments. Using the ﬁrst three principles plus that of crosscutting relationships, he inferred two marine advances of the sea separated by a period of tilting and erosion: what we would call an unconformity. Steno thought the younger strata were deposited from Noah’s ﬂood, and the older set in the universal ocean after the second day of creation. His diagram of strange collapsed slabs is probably due to the local geology of Tuscany, which consists of deformed strata, including large synsedimentary slides, eroded into sharp-edged peaks (Fig. 6.3). Steno’s revolutionary ideas still form the basis of stratigraphy. They were elaborated on during the 18th century and were a key factor in the cyclical theories of Hutton. Unfortunately Steno’s principles can only be used to correlate strata in small areas. Strata comparable to those around Florence, although showing similar relationships, may be very different in age in distant areas. For example, the various delta lobes of the modern Mississippi River are of different ages at different scales, even though individual lobes may show similar stratigraphic sections (Figs 5.9 and 6.4).

(b)

Fig. 6.3 (a) Steno’s diagram of stratigraphic relations in Tuscany showing repetitive (25 to 20) deposition–solution–collapse changes through time (from Gohau 1991, ﬁg. 25); (b) isolated tilted strata, Valdorbia, Umbria.

Fig. 6.4 Mississippi subdeltas of different ages forming distinct lenses with dates of initial crevassing indicated (from Anstey & Chase 1979, ﬁg. 14.7A,B).

118 Chapter 6

Fig. 6.5 Restored cross-section of the Cambrian rocks in the Grand Canyon (from Petersen & Rigby 1990, ﬁg. 5.4). (Copyright 1990. Reproduced with permission of The McGraw-Hill Companies.)

Methods are needed that let you to equate separate rock layers, or determine their relative ages, in separate areas. The ﬁrst useful method discovered was based on characteristic fossils (see Chapter 7). However, within one sedimentary basin, it is possible to physically correlate strata using various methods. 6.1.1 Lateral tracing of strata Simple lateral tracing of beds is possible in exceptional circumstances, such as on unbroken faces of canyon walls, where individual beds and bedding planes can be traced by eye, or by walking along the continuous outcrops. In practice, however, even in canyons, the beds usually die out or pass into other rock types eventually. In the ﬁnest example of physical continuity, the walls of the Grand Canyon of the Colorado River, and the standard unit from that canyon cited as the ﬁnest example, the Tapeats Sandstone, the bedding planes are impossible to trace far before they split, amalgamate, die out, or otherwise disappear for a variety of (usually indeterminable) reasons. In practice, even in the Grand Canyon, sections

were measured some distance apart and lithological similarity used to connect them (Fig. 6.5). 6.1.2 Seismic cross-sections Seismic cross-sections look like tracings of the layers in a well-exposed canyon wall (Fig. 6.6). But this is misleading. Corrections and assumptions must be made before the raw seismic data can be plotted and interpreted. Anyone using seismic methods in stratigraphy should be aware of the difﬁculties involved and the corrections required to try to overcome them, or they had better not use such records (Sheriff 1980; Anstey 1982; Sheriff & Geldart 1982a,b). SEISMIC STRATIGRAPHY essentially involves making big bangs at the earth’s surface and listening to the sound echoes reﬂected back from underground surfaces (Sengbush 1983). As with echoes in the atmosphere, the clarity of subsurface echoes depends on how sharp and planar the reﬂector is: shouting across a canyon with a sheer and ﬂat face will give you a good echo; shouting at a line of trees will not.

The horizontal dimension 119

Fig. 6.6 Seismic cross-section, offshore Japan. The water-bottom reﬂection, A, with travel time 1.0–2.0 s, indicates water depths of 750–1500 m. The ship travelled 8.5 km between the 30-minute marks at the top. Most primary reﬂections are obscured by the ﬁrst and second water-bottom multiples, B and H. More than a kilometer of sediments are indicated by the reﬂections near C; multiples of these appear paged-back at D. E indicates a fault scarp on the ocean ﬂoor. F shows diffractions, probably from sea-ﬂoor relief slightly offset from the line. Note the onlap and thinning above G (from Sheriff & Geldart 1982a, ﬁg. 5.56). Reproduced with permission of Cambridge University Press.)

Seismic waves travel though rocks at different speeds depending on seismic velocity and rock density. So, it is possible to infer rock density from seismic waves of known initial velocity in a rock of known density. This is how we infer rock densities deep within the earth. Also, seismic waves are reﬂected at junctions between different rocks where the acoustic impedance changes. ACOUSTIC IMPEDANCE is the product of seismic velocity and rock density. Thus, a shale–sandstone junction may show up as a well-marked reﬂector in a seismic section, and such surfaces may be traced across a section to show its basic structure.

Seismic sections are produced by transmitting artiﬁcially produced sound waves into the earth and recording and analyzing the reﬂections from surfaces underground. The dominant method of reﬂection seismic analysis, whereby reﬂections from surfaces are analysed, has now mostly supplanted the older refraction seismic studies. On land, sound waves are produced either by exploding dynamite in shot holes or by dropping heavy weights onto the ground, and the reﬂections are then picked up by arrays of geophones laid on the ground. On water, sound waves are produced by airguns and the reﬂections are picked up by ﬂoating arrays of

120 Chapter 6

Fig. 6.7 Offshore seismic methods (from Press & Siever 1986, box 18.1).

hydrophones (Fig. 6.7). In both cases, the geo- or hydrophones are arranged in various arrays designed to minimize interference. After ampliﬁcation and ﬁltering to reduce noise, the signals are transmitted and recorded on magnetic tape in a recording truck. In the days of paper records, geophysicists then had to grapple with static and normal moveout corrections on a blizzard of paper records before a sensible interpretation could be made. Personal decisions about these corrections affected the result. Now almost anyone can look at a seismic section and make a casual interpretation without realizing that someone has already decided what assumptions were made in the computerized corrections (Bally 1987). After computer processing, the data for each geo- or hydrophone are plotted on a distance versus time section. The traces of individual phones are often close enough so that the higher amplitude deﬂections (wavelets) overlap and characteristic lines can be traced across the record section (Fig. 6.6). Now comes the need to understand what you are looking at: certainly not a pristine plot of rock layers and structures. You must take into account signal noise, signal attenuation, lack of acoustic contrast, and signal interference. First, only a weak signal returns after reﬂection from a depth of a few thousand meters, a signal obscured by random noise. A correction to delete random noise is needed. Second, high frequencies have only low penetration due to energy loss. Only low frequencies survive pene-

tration through a few thousand meters of rock. Frequencies change from perhaps 100 Hz near the surface to 30 Hz at several thousand meters. This roughly translates into respective wavelengths of around 40–100 m. The wavelets on the seismic plot obviously do not mark sharp lithological boundaries and encompass a typically entire road section of several lithologies (Fig. 6.8). Third, not all layers of interest show up. The strength of the reﬂection depends on acoustic contrast across rock junctions. Sharp lithological changes across horizontal surfaces usually produce the most marked reﬂectors. Irregular and inclined surfaces and gradational lithological boundaries may not give coherent reﬂectors, particularly if acoustic impedance does not change much. Thus internal bedding surfaces within homogenous lithologies will not show up, and the structure of, for example, limestone sequences may not be determinable. In the Auk Field of the North Sea, the critical contact between the Zechstein carbonate reservoir and the overlying Upper Cretaceous Chalk simply does not show on the seismic sections. Furthermore, actual lithologies cannot be determined directly, which is why test boreholes are necessary in order to calibrate the reﬂector horizons with actual surfaces and rocks. Fourth, at times longer than about 4 s, interference between returning signals may completely obscure the reﬂectors, or spurious reﬂectors may appear. Primary seismic reﬂections are caused by density/velocity contrasts, so that water–sediment, oil–water, gas–water

The horizontal dimension 121

Fig. 6.8 Scale of typical seismic wave against large outcrop, and well log (on left) (based on Miall 1990, ﬁg. 1.6 (Copyright (1990). Reproduced with permission of Springer-Verlag) and Anstey 1982, ﬁg. 3.3).

contacts also cause reﬂections and interference may duplicate such surfaces. Common misleading surfaces in offshore studies are multiple sea-bed reﬂectors (Fig. 6.6). Fifth, the velocity of sound varies laterally and vertically within the earth. So it is hard to convert the time sections into true depth sections, which is one reason why most sections are shown as distance against time. Sixth, inclined layers give multiple reﬂections that can interfere to give spurious reﬂectors, as in the case of faults. These problems are partly overcome by corrections made during processing. However, the choice of corrections made during computer processing requires assumptions about the lithologies, attitude, and structure of the area being studied. For example, inaccurate assumptions about structure may produce simple gentle folds in areas of sharply compressed and overturned folding. The following are typical corrections. STATIC CORRECTIONS overcome the effects of changes in elevation at source and receiver locations, changes in source and receiver depths, variations in velocity and

thickness of weathered zones on land, and variations of water depth and sub-bottom velocity on marine data. Static corrections normalize the signal to a datum. Dynamic corrections overcome the effects of offset distances between source and receiver for both horizontal and dipping beds. The correction (normal moveout) can be complex but basically it transforms the actual distance from source to reﬂector to a vertical distance from the datum. Deconvolution is an attempt to overcome the severe distortions due to multiple reﬂections (ghosts and reverberations) in near-surface seismic sections. In processing, deconvolution is normally applied before normal moveout corrections because normal moveout differentially stretches the data and causes spectral changes. Horizontal stacking adds the reﬂections from different traces after normal moveout corrections. It enhances the primary reﬂector by further removing residual normal moveouts and further suppressing random noise.

122 Chapter 6

Fig. 6.9 Seismic sections: (a) before migration; (b) after migration (from Sengbush 1983, ﬁg. 1.6).

Migration repositions reﬂections at their originating surfaces, compresses diffractions back to their originating points, and moves multibranches into their true structural position. Migration is required because all primary seismic data are orientated with respect to the recording instruments. On a time section before migration, faults produce diffractions, synclines produce multibranch reﬂections, and anticlines appear broader than they actually are (Fig. 6.9). You get a truer picture of the structures below the surface after migration. This was recognized right at the beginning of seismic exploration: the very ﬁrst seismic reﬂection data in 1921 were migrated by hand! Lack of attention to these corrections or mistakes in applying them are the main reasons why data become useless, often without anyone realizing it (Sengbush 1983). Despite their apparent precision and all the computerization involved, the usefulness of seismic studies depends on the interpretative powers of the geophysicists and geologists making them. These can only be acquired by practice in translating the results of seismic experiments into geologically valid models of earth structure (McQuillan et al. 1984). Interpreting seismic sections accurately is therefore something of an art, requiring long experience and a thorough knowledge of geologic setting. Seismic sections give little information on rock type, porosity, permeability, or many of the other properties we are interested in, and must be calibrated with data from natural sections or boreholes (Fig. 6.10). 6.1.3 Lithological similarity Equating similar rock units or successions in separate OUTCROPS or boreholes is the ﬁrst stage in physical corre-

lation, if continuous outcrops or seismic sections are not available. 6.1.3.1 Outcrops In all but the most boring natural sections there will be some change in rock type upwards, either a sharp or a gradual change. By now, you will (hopefully) be able to deduce the processes and depositional environments represented by the sediments and can hence choose reasonably signiﬁcant natural boundaries between rock units. With luck, the same boundaries will turn up in adjacent local sections and perhaps they will also show up topographically between exposed sections due to the different weathering characteristics of the adjacent units. The boundaries can then be traced on a map to show the disposition of the rock units. Rock units will frequently: (i) disappear; (ii) increase or decrease in thickness; (iii) change in character laterally; and (iv) become interbedded with other units, as shown in Fig. 5.18. These rock units can then be given informal or formal stratigraphic names to aid description, and particularly good local “marker” horizons may be individually named. In many regions, rock divisions are harder to make. Often this is not because of uniformity but because of variety. For example, the Ordovician carbonates of eastern North America have monotonous repetitions of endlessly variable lithologies interbedded in apparently bewildering and haphazard ways. For almost a century, the rocks have deﬁed satisfactory subdivision. The basic mapping unit is a rock unit whose only rationale is that it can be physically traced over the (frequently vegetation-covered) landscape. In Europe, the names for these units used to be informal. In North

The horizontal dimension 123

Fig. 6.10 Correlation of seismic horizons and well log (from McQuillan et al. 1984, ﬁg. 10.7). (With kind permission of Kluwer Academic Publishers.)

America, such units are called formations and formally named by type locality and lithology (e.g. Navajo Sandstone) or simply by type locality and formation if there is no characteristic or dominant lithology (e.g. Kayenta Formation). However, in only a few remote areas can you start from scratch (Fig. 6.11). Usually, other people have already deﬁned, redeﬁned, and argued over the stratigraphic names of an area; so you must do the best you can with the result (see Chapter 9).

(bailed out) during drilling. Being rather coarse, these CHIPS gave good information on lithology, porosity, and the presence of oil or minerals. Modern rotary drilling, however, gives only ﬁne cuttings contaminated with drilling ﬂuid and up-hole caving of the walls. Solid BOREHOLE CORES are more reliable but coring boreholes is expensive and is now only used to calibrate indirect information from instrument logs. BOREHOLE

6.1.3.3 Wireline logs 6.1.3.2 Borehole chips and cores In the early days of drilling, when exploration wells were cut with cable tools, cuttings were removed with pails

Instrument logs (or WIRELINE LOGS) involve lowering an instrument down the borehole on a cable and measuring various properties of the borehole wall as it

124 Chapter 6

Fig. 6.11 Correlation of Upper Devonian sections in Montana, showing use of informal units during reconnaissance studies. Standard symbols used except where indicated (from Krumbein & Sloss 1963, ﬁg. 3.8). (© 1951, 1963 by W.C. Krumbein and L.L. Sloss. Used with the permission of W.H. Freeman and Company.)

descends. Different logs give different information on the rock type, dip of the strata, porosity, fracturing, water salinity, and hydrocarbon content. The value of such logs depends on knowing how various tools respond, understanding the effects on the rocks of drilling the borehole, and integrating the logging and log interpretation with other geological data (Alger 1980; Luthi 2000). Early electrical logs recorded two types of curves: spontaneous potential, which is mainly a sandstone–shale discriminator; and resistivity curves, which are of many types. SPONTANEOUS POTENTIAL (SP) logs are determined by several factors. In thick porous formations, the SP deﬂection from the shale baseline is due to the contrast between the salinity of the formation waters and the borehole ﬂuid. With low-salinity drilling ﬂuids, the SP deﬂects to the left opposite formations containing saline porewaters, such as sandstones. With saline

drilling ﬂuids, there is no deﬂection or only a small deﬂection. RESISTIVITY (R) LOGS measure the resistance of the borehole walls to electric currents of different amperages and voltages. Deﬂections on the resistivity logs can then be calibrated to known formation lithologies. These two logs are often the ones given in stratigraphic exercises (Fig. 6.12). Porosity logs respond predictably to variations in porosity, contained ﬂuids and lithology, and consist of sonic, formation density, and neutron logs. Sonic logs measure the transit time of an ultrasonic compressional wave over a standard length of borehole (Fig. 6.13). Formation density logs measure the bulk density from the ﬂux of applied gamma ray radiation to the borehole walls, and the porosity is calculated by comparing the grain and cement densities with the bulk density. Neutron logs basically measure the hydrogen ion concentration. High hydrogen ion concentrations can be due to

The horizontal dimension 125

Fig. 6.12 Spontaneous potential and resistivity logs of a borehole.

good porosity ﬁlled with oil or water, or to shale. Porosity evaluation is based on response to known lithologies and porosities in test pits. Natural gamma ray logs measure the gamma ray ﬂux from the formation due to decay of radioactive potassium, uranium, and thorium (Fig. 6.13). Shales usually have the highest abundances of these elements. So natural gamma ray logs can substitute for spontaneous potential logs, where SP logs are not available because of an empty hole, oil-based mud, saline mud, or a cased hole. Dip-meter logs measure the inclination of layers and are useful in sedimentological studies, particularly of large-scale cross-bedded units such as eolian sands (Fig. 6.13).

All these logs obviously have problems of interpretation due to similar responses to a variety of causes. Nevertheless, geologists could see characteristic patterns repeating from borehole to borehole even with the earliest logs. These patterns correlated with variations in shale, sandstone–shale, and calcareous sequences. Physical correlations could be made between boreholes that were much more deﬁnite than could be obtained from well cuttings or drilling times. In fact, the main use of wireline logs has been for physical correlation. Both spontaneous potential and resistivity logs allow excellent bed identiﬁcation and allow various structures such as folds and faults to be easily determined from closely spaced boreholes (Fig. 6.14).

126 Chapter 6

Fig. 6.13 Gamma-ray, sonic, and dip-meter logs in Permian sandstone (from Glennie 1982, ﬁg. 3). (Copyright (1982) with permission from Elsevier Science.)

6.1.3.4 Unique layers Sometimes unique rock layers can be identiﬁed and traced throughout an area in natural exposures or boreholes. These layers are often caused by rare, short-lived events and hence give not only physical but also time correlation. The events can be explosive volcanic eruptions, extraterrestrial impacts, intense storms, or any other widespread unusual natural event. The main problem is in demonstrating that the layer is in fact unique, and not simply two or more distinct eruptions, impacts, or storms variously preserved in separate sections. In the ﬁrst part of the 20th century, bentonite correlation was widely applied to the Lower Paleozoic

sections of eastern North America. However, the occurrence of bentonite in any section is determined both by subsequent preservation and by the direction of the prevailing wind during the eruption. Ash beds can be absent from some sections if they are removed by erosion before overlying beds are deposited. Successive ash falls rarely lie in nice circles around a volcano, but in often non-overlapping ellipses (Fig. 6.15). Preservation of the individual layers was thus too haphazard for correlation, particularly since the individual ash layers from eruptions could not be satisfactorily distinguished. More recently, geochemical ﬁngerprinting of individual ashes has allowed some correlations to be made (Fig. 6.16). However, variations in

The horizontal dimension 127

Fig. 6.14 Surface–subsurface rock unit correlation with spontaneous potential and resistivity logs (from Krumbein & Sloss 1963, ﬁg. 10.4). (© 1951, 1963 by W.C. Krumbein and L.L. Sloss. Used with the permission of W.H. Freeman and Company.)

geochemistry can be due to separation of minerals by size and density during eruptions, as well as to secondary causes, so some problems remain (Huff et al. 1996).

6.2 Lateral changes Recent sediments change character laterally, controlled by the extent and variability of the Recent environments in which they formed. Thus, contemporary alluvial fan deposits have rapid and frequently abrupt transitions between coarse and ﬁne sediments over short distances, while those of abyssal plains show monotonous expanses of uniform clays over wide areas.

Ancient sedimentary rocks must change laterally in the same way. A FACIES is simply a distinct variety of rock or rocks and/or their contained fossils; FACIES CHANGE simply refers to the change in rock type or fauna laterally. However, lateral changes can be difﬁcult to prove in practice. Philip B. King (1949) described his discovery of the classic reef to back-reef facies change in the Permian Capitan Reef of Texas as follows: “A blackboard drawing or a textbook illustration of a sedimentary facies has quite a different appearance from a sedimentary facies when one encounters it in the ﬁeld. In the ﬁeld, facies changes seem bafﬂing and bewildering, especially in an area of new or unknown stratigraphy . . . I remember vividly, for

128 Chapter 6 (a)

Fig. 6.15 (a) Plume from Rabaul volcano; (b) ashfall ellipses from Mount St Helens 1980 eruption.

example, a summer afternoon in the Sierra Diable in 1928, when we climbed a projecting angle of a canyon wall and saw spread before us the stratigraphy of the Permian rocks of the interior of the range. In the preceding weeks we had been tracing out the stratigraphy of the Permian on the frontal escarpment of the range, and had divided the rocks into four or ﬁve wellmarked lithologic and faunal units of limestone, dolomite, and shale. But now, as we looked into the interior of the range, we

saw all our ﬁne units dissolve before our eyes, merging into a monotonous sequence of thin-bedded dolomite that extended as far up the canyon as we could see.”

In many areas, consistent rock sequences with their characteristic fossil biotas demonstrably persist over large areas. Faced with these apparently uniform

The horizontal dimension 129

Fig. 6.16 Geochemical bentonite correlations for Ordovician (from Kolata et al. 1986, ﬁg. 15). (Reproduced with permission of the Illinois State Geological Society.)

successions, workers have often interpreted them in classical Wernerian terms as caused by successive changes of environments through time in a practically universal ocean. The eventual change into other rocks has often come as a surprise. 6.2.1 Facies concepts The term “facies” simply means aspect, look, or appearance. It is normally used to refer to the total rock and fossil characteristics of a stratigraphic unit that set it apart from other adjacent units (Hallam 1981). However, the concept has often given geologists a lot of trouble. Furthermore, the basic question debated by Werner and Hutton still needs asking in any stratigraphic study: are the vertical changes in rocks and fossils shown by a succession due to changes of environment through time, or are they due to the lateral migration of environments during sedimentation? Or as Ager (1973) put it, are sediments deposited by the “gentle rain from heaven” or by the “moving ﬁnger [which] having writ moves on.”

During the 18th and early 19th centuries, the most obvious feature of strata appeared to be their lateral persistence. Similar distinctive successions, such as the Cretaceous Greensand–Gault–Chalk succession, could be traced across Western Europe. This led to Werner’s Universal Ocean, to William Smith’s Law of Faunal Succession, and to Cuvier’s catastrophism. Sediments were deposited and faunas were created in a universal ocean, and were then overwhelmed and buried, only to be replaced by other sediments and more complex faunas in a new universal ocean. Successive ﬂoods and faunas in a sequence of universal oceans deﬁned the geological systems. This approach still survives in a highly modiﬁed form, because it sometimes ﬁts the rock and fossil record. Furthermore, large meteorite impacts have undoubtedly caused biological catastrophes and redirected evolution at least once during the Phanerozoic (Alvarez et al. 1980). However, as catastrophism was being developed, people studying modern environments realized that neither contemporary sediments nor contemporary organisms

130 Chapter 6 (a)

Fig. 6.17 (a) Chateau du Joux, disrupted anticline of Upper Jurassic limestone (courtesy of Interfrance); (b) present and restored cross-sections of anticlines (from Bitterli 1990, ﬁg. 6).

were uniformly distributed over the earth’s surface. Both sediments and organisms varied markedly across the globe, depending on local environmental conditions. Also, similar sediments and organisms recurred whenever and wherever local environmental conditions were suitable. The uniformitarian counterconception to catastrophism, developed by Hutton, was expanded during the early 19th century. The adoption of Lyell’s strict uniformitarian approach to ancient rocks soon led to the development of the facies concept by Gressly in 1838 (Nelson 1985). The conﬂict between universal ocean and facies change persists throughout the history of stratigraphy; for example, in the controversies of sequence stratigraphy – universal oceans (eustatic sea-level changes) versus facies changes (environmental changes). It is a fundamental problem in all historical studies – “time’s arrow versus time’s cycle,” as Gould (1987) put it. Yet, we need both concepts: the arrow of historical unique-

ness (universal oceans) to explain changes caused by the irreversible chemical, physical, and biological evolution of the earth, and the cycles of unchanging processes (facies changes) to explain the repetition of rock and fossil types in sections. Gressly developed his facies concept in the Upper Jurassic deposits of the Swiss Jura mountains around Solothurn (the same area where Oppel (1856–8) later developed the zone concept). Here, the Jura mountains are deformed into box folds and there is a marked change in rock types across strike from one faulted anticline to another, since the intervening rocks have been eroded away (Fig. 6.17). Gressly speciﬁcally included both lithological and biological aspects in his somewhat involved High German deﬁnition: “there are two principal points that always characterize the group of modiﬁcations that I call facies . . . one is that a similar petrographic aspect of any unit necessarily implies, wherever it is found, the same pale-

The horizontal dimension 131

Fig. 6.18 Swiss Jura during shifting depocentres, Oxfordian: (a) end of the tranversarium zone; (b) end of the bimammatum zone; (c) end of the planula zone (from Gygi 1986, ﬁg. 3).

ontological assemblage; the other, that a similar paleontological assemblage rigorously excludes the genera and species of fossils frequent in other facies” (Gressly, translated by Middleton 1978). What this means is that different contemporary rock types (lithofacies) always contain different fossils (biofacies). This extreme view was based on Gressly’s spectacular environmental changes in the Middle Oxfordian, from reef limestones to deep-water calcareous shales in less than 5 km. However, because these changes occurred across the eroded Jura anticlines, Gressly could not prove many of his facies transitions (especially since Oppel’s later reﬁned zone concept only applied to the deeper water shales). This leads to a curious irony: Gressly based his facies

concept on a succession that does not show facies changes. The Middle Oxfordian platform basin succession, on which Gressly based his concept, consists of overlapping lenses or sequences which formed parallel to the Jura folds and which succeeded each other in time (Fig. 6.18a,b). Only in the Upper Oxfordian is there a contemporary reef-basin facies change (Fig. 6.18c). Facies develop because of variations in sedimentary environments. So, a coarse beach gravel or sand might change offshore into quieter water silts, clays, and ﬁnally perhaps carbonate mud. The earth has a variety of sediment being deposited at any one time, as well as a variety of organisms living on and in these sediments. The same organisms may occur in the same sort of sediment,

132 Chapter 6

Fig. 6.19 Facies relationships: (a) with uniform sedimentation and static environmental boundaries; (b) with uniform sedimentation and migration of facies boundaries over time. The heavy time-lines are horizontal and equally spaced (from Petersen & Rigby 1990, ﬁgs 5.1, 5.2). (Copyright 1990. Reproduced with permission of The McGraw-Hill Companies.)

in which case the sedimentary and organic distributions correspond. However, organisms may be controlled by factors unrecorded in the sediment (such as salinity and temperature). So, in some cases, the distribution of organisms is different from the distribution of sediments. For this reason, we separate LITHOFACIES, based on sediments (or sedimentary rocks), from BIOFACIES, based on organisms (or fossils). Sometimes they correspond; often they do not. 6.2.2 Lithofacies If the boundaries between lithofacies remain constant as sediment builds up, then a series of vertical facies belts result (Fig. 6.19a). For this to occur, depth of water, rate of sedimentation, composition and texture of introduced sediment, rate of subsidence, and a host of others factors, have to remain constant. This is practically unknown. Usually the facies boundaries migrate over time, as sediment builds up, causing the belts and boundaries to shift laterally with time (Fig. 6.19b). In the example shown in Fig. 6.19(b), facies boundaries have shifted to the left as clastic sediment is supplied from the right. This might be caused by a drop in sea level, a decrease in tectonic subsidence, or simply the natural build out of sediment from the land. Even this diagram is unnatural: the time-lines are parallel and equally spaced, so sediment is

accumulating at constant rates all over the area, even in different environments – a very unlikely situation. The Grand Canyon example of lateral tracing is also a good example of facies change. The two dashed lines in Fig. 6.5 are time-lines based on fossils. From this you can infer that the facies boundaries between the Tapeats Sandstone, Bright Angel Shale, and Muav Limestone become younger to the east (a mirror image of Fig. 6.19b). At any one place the facies change is from nearshore to more offshore environments. However, differential rates of sedimentation usually give divergent time-lines. In Fig. 6.20, the rate of sedimentation, and probably subsidence, was more rapid on the right of the diagram, and thicker sediments accumulated. This is the most natural of the three crosssections. In the Devonian Catskill Delta of New York State, black transgressive shales extend eastward from the basin to separate regressive delta build-out from the rising Appalachian Mountains (Fig. 6.21). The transgressive black shales are almost time-lines and show that maximum subsidence occurred in the facies change from red beds to marine sands. Facies studies can therefore provide information on relatives rates of sedimentation, direction of transport, and capacity of the transporting medium. Both the Devonian Catskill Delta and the Cambrian Grand Canyon sections demonstrate WALTHER’S LAW (or

The horizontal dimension 133 Principle), which states that in a continuous vertical sequence, the succession of facies vertically reﬂects the original contemporary lateral arrangement of facies in the area (Fig. 6.22). 6.2.3 Biofacies The assemblage of fossils in sediment can to help environmental and paleogeographic interpretations as well as giving the relative ages of rocks (Chapter 7). All organisms have particular modes of life and speciﬁc tolerances, which are controlled by a host of environmental factors, many of which are not preservable in sediment. If these modes of life and tolerances can be determined for fossils, then additional information on factors not preserved in sediment (such as turbidity, salinity, oxygen content) can be inferred (Fig. 6.23).

Fig. 6.20 Facies relationships produced by varying subsidence and/or sediment supply. Time-lines diverge where the net sediment rate increases (from Petersen & Rigby 1990, ﬁg. 5.3). (Copyright 1990. Reproduced with permission of The McGraw-Hill Companies)

Simple comparisons may be misleading because of evolutionary changes through time. For example, most living brachiopods are ﬁxed, immobile suspension feeders, but many fossil brachiopods had special adaptations for free-living. The living grizzly and polar bears are anatomically almost identical, but have very different modes of life. Nevertheless, biofacies can amplify and reﬁne environments based on lithofacies, and allow more detailed analysis of environmental variables (Fig. 6.24). Biofacies may be useful in inferring hydrographic conditions in some environments. For example, in the Devonian deltaic succession of New York State, the Middle Devonian Sonyea Group has two deltaic units separated by black shales (Fig. 6.25). The lower delta has sandy shales grading into shales with no barrier, and resembles the pro-delta–delta front transition of actively prograding deltas such as the Mississippi Birdfoot Delta (Fig. 4.21). Here the biofacies (Productella and Bellerophon communities) do not correspond with the lithofacies and are very dissimilar. The Bellerophon community has many more saline-tolerant molluscs and very few brachiopods, suggesting a lower-salinity environment. The biofacies boundary was probably controlled by the divide between an upper low-salinity wedge from the distributaries and underlying normalsalinity marine water occupying most of the pro-delta slope (Fig. 6.26a). The upper delta has a sandy barrier-bar unit and resembles the reworked barrier–lagoon complex of the Mississippi (Fig. 4.25). Here, the biofacies (Rhipidomella and Cypricardella communities) correspond with the lithofacies, and are similarly dominated by brachiopods.

Fig. 6.21 Restored cross-section at the close of Devonian sedimentation in New York (from Kay & Colbert 1965, ﬁg. 11.9). (Copyright © 1965. This material is used by permission of John Wiley & Sons, Inc.)

134 Chapter 6 Both biofacies and lithofacies are controlled by the physical features of the environment, since the breaking waves and reduced freshwater input provide nearnormal marine salinity environments (Fig. 6.26b).

6.3 Mapping

Fig. 6.22 Walther’s Principle: ideal vertical facies change with transgression and regression.

Maps can show the distribution not only of rocks, but of many other useful features of sediments such as thickness, facies, current direction, composition, and fossil zones (Barnes 1981). However, units are rarely completely exposed and a geologist rarely has time to examine the entire area in great detail. So, most maps are based on limited information, and extrapolations have to be made between sample control points (see Tearpock & Bischke 2001). What is actually visible in the ﬁeld is placed on outcrop maps (Fig. 6.27) From these, geologists then construct interpretative geologic maps

Fig. 6.23 Response of certain organisms to salinity (from Anstey & Chase 1979, ﬁg. 5.2).

Fig. 6.24 Lower Cretaceous shallow marine to brackish clastic communities (biofacies) in relation to substrate (lithofacies) (from Scott 1974, ﬁg. 11).

Fig. 6.25 Sonyea Group with biofacies and lithofacies plotted against time (not thickness!) (modiﬁed from Sutton & McGhee 1985, ﬁgs 2, 3). (Copyright © 1985. Reproduced with permission of the publisher, the Geological Society of America, Boulder, Colorado, USA.)

136 Chapter 6

Fig. 6.26 Reconstructions of deltas of the Sonyea Group (from McGhee 1976, ﬁg. 16). (a) Mississippi-type delta with rapidly prograding low-salinity distributaries. (b) Nile- or Mississippi-type barrier–lagoon with mixed near-normal salinity water almost to shoreline.

Fig. 6.27 An outcrop map showing observed sedimentary rocks and attitudes (modiﬁed from Bennison 1969, map 20). (Reproduced by permission of Hodder Arnold.)

The horizontal dimension 137

Fig. 6.28 An interpretive map constructed from the outcrop map shown in Fig. 6.27, completed with standard geologic map extrapolation techniques.

(Fig. 6.28). A combined outcrop/interpretative map shows how much extrapolation went into drawing the interpretative map. Unfortunately, many maps give only the interpretation, so it is important to do some mapping yourself; otherwise you may have an unwarranted conﬁdence in maps done by others. Many stratigraphic type sections are based on exposures even worse than those shown in Fig. 6.27, which is not obvious from the interpretative maps used to locate them. For example, to examine the type section of the Ordovician Caradocian Series, you wade down the muddy Onny River in Shropshire, England, trying to hammer samples out of the (frequently) gravel-covered stream bed. Map reliability and resolution should be of vital concern to all geoscientists, yet these factors are usually ignored in map-based studies. Most studies have too few data points to adequately resolve complex

environments by mapping, even where time equivalence of the points can be assumed. And beyond the Quaternary time equivalence is rarely resolvable within several million years. Dennison (1972) showed the dramatic effect of sample density on environmental resolution, even where the samples were deﬁnitely time equivalent (Fig. 6.29). Section descriptions suffer from the same problem in that section-thickness precision varies by over 10% at the 95% conﬁdence level. This affects, for example, the reliability of isopach maps that show thickness variations of units (see Fig. 8.4). In fact, as Dennison noted, you should know the density of samples required to attain your mapping objectives before you start the mapping. Or you should understand the limits that existing outcrops, sample densities, and dating precision place on interpretation. Statistical analysis of signiﬁcance

Fig. 6.29 Aeronautical chart of the Mississippi Delta, showing how the resolution of the three main environments improves as the number of random sample points increases (from Dennison 1972, ﬁg. 1). (Copyright © 1972. Reproduced with permission of the publisher, the Geological Society of America, Boulder, Colorado, USA.)

The horizontal dimension 139 should be done before, during, and after sampling. On the one hand, the size of the database determines the precision possible; on the other hand, the precision required determines the size and distribution of the database required. Fanciful reconstructions based on limited databases should be discarded.

Isopach, current direction, facies, paleotectonic, and other maps can show basin evolution through time. These maps are discussed in the basin analysis chapter (Chapter 8).

7 The time dimension

7.1 7.2 7.3 7.4

Age equivalence Relative ages Numerical methods (ages in years) Calibration of relative and numerical dates

“A people without history is like the wind on the buffalo grass.” Traditional Sioux saying

Geologic time is unimaginably long, linear, and continuous, beginning (for the solar system) with the formation of the sun and surrounding planets about ﬁve billion years ago. This immensity of time is geology’s fundamental contribution to Western thought (Gould 1987). The scale of geologic time was discovered only relatively recently and its linearity and continuity were not always accepted (Eicher 1976). Theories of the earth’s formation are as old as human history; but most were qualitative and neither linear nor continuous. Aristotle (384–322 bc) observed the perfect circular motions of the planets and thought that time must be perfect and circular as well. Stoic philosophers elaborated this into the idea that everything that has ever happened will happen again in endless cycles through eternity. The ancient Mayas thought that time was linear but interrupted by periodic stops and starts

(Albritton 1984). The Hindu Vedic scriptures come closest to our modern concepts. Each cycle, or Great Year, was about four million years long. One thousand Great Years (four billion years) make up one kalpa, or one day, in the life of the creator Brahma (which is pretty close to the current estimate of the age of the earth). One hundred kalpas (400 billion years) are the life-span of Brahma, after which the earth will dissolve for about 3000 billion years, and then will be reborn (Wagar 1982). However, none of these ideas were based on natural observations. The idea that time is linear and continuous comes to Western science from the early Christian fathers. St Augustine (ad 354–430) insisted that Christ could die for man only once. So, the Cruciﬁxion had to be a single event, not an endless repetition on the wheel of time; otherwise history would make a mockery of religion. History therefore must be a series of unique events with a beginning at the Creation and an end at the Last Judgement. This prompted various divines to attempt to date the creation from analysis of scripture. (We do the same thing today, except that our analysis is of the rock, not

The time dimension 141

Fig. 7.1 Age estimates of the earth, ad 1650 to the present day.

the human, record.) Between ad 100 and 300, various estimates clustered around 5500 bc. The most precise (and most famous) estimate of 4004 bc was made in the 17th century by Bishop Ussher of Ireland, based on Biblical genealogies. These short estimates were usually accepted in the 17th and 18th centuries. Natural philosophers asked not how long it had taken for geologic forces to mold the earth, but how they had managed to do so much in the few millennia since the Creation. One big problem with the short estimates was evidence that the earth had changed many times since the Creation. By the late 18th century, it became increasingly difﬁcult to compress rock sequences with their varied fossils, and thus the history of the earth, into a few thousand years. Both geologic studies and advances in physics expanded the short estimates (Fig. 7.1). Buffon liberated earth history from human timescales by extending the origin of the earth to some 75,000 years, based on estimates of the earth’s rate of cooling. Hutton expanded this further to an endless “no vestige of a beginning, no prospect of an end.” By the mid-19th century, with the publication of Lyell’s Principles of Geology in 1833, and Darwin’s Origin of Species in 1859, both geology and biology needed immense periods of time to explain the observed changes in rocks and fossils. However, there was no way of calculating exactly how much time these processes needed, since both geologic and biologic rates varied enormously, and estimates based on physics were orders of magnitude less than the time most biologists and geologists required for their histories. Rough geologic estimates of the minimum age of the earth had been made since the 17th century. For

example, the astronomer Halley estimated the age of the oceans by dividing the total amount of salt in the sea by the rate of salt supply by rivers. Other estimates were made by dividing the total thickness of the stratigraphic column by the average rate of sedimentation. However, rates of salt supply to the sea and rates of sedimentation are variable. Salt is also removed from the sea in porewaters, in reactions with submarine basalts, and in salt basins, and sedimentary rocks are eroded and redeposited during Hutton’s repetitive rock cycles. Nevertheless, by the late 19th century, age estimates using these methods gave 90 to 100 million years for the oceans (determined by Joly in 1899), and about the same for sediments (determined by Reade in 1879) (Burchﬁeld 1975). These estimates were really only accepted because they ﬁtted the mid-19th century estimates based on physics (see below). In fact, both methods give enormously different values depending on the assumptions made: anywhere between 1500 million years and 28 million years. The unlimited time demanded by uniformitarians in the ﬁrst half of the 19th century clashed head-on with the laws of physics as then understood. Buffon’s earlier work was expanded and developed in the later 19th century. Lord Kelvin calculated the age of the earth to be less than 100 million years in 1862, and ﬁnally proposed a mere 24 million years in 1897. Geologists and biologists made a ﬁght of it, but by 1897, virtually no-one thought in terms of hundreds of millions of years for earth history (Burchﬁeld 1975). Ironically, Kelvin’s ﬁnal word was given shortly after the discovery of radioactivity by Becquerel in 1896. Radioactive decay provided another source of heat, which not only demolished all Kelvin’s careful calcula-

142 Chapter 7 tions, but also provided for the ﬁrst time an accurate way of dating minerals and rocks. Experimenters discovered that various radioactive elements decayed at a constant rate, and hence the total decay products accumulated over time could be compared with the rate of decay to give an accurate date for the formation of the mineral or rock. The only theoretical limitation was the rate of decay of the radioactive element: too fast a rate of decay meant that only very recent events could be dated; too slow a rate meant that not enough decay products accumulated. By 1907, Boltwood had produced U-Pb dates from zircons ranging from 400 million years old to an incredible 2200 million years old (Burchﬁeld 1975). Geologists were at ﬁrst reluctant to accept the new ages. Throughout the 19th century, physicists had insisted on a short (less than 100 million years) age for the earth; now they were insisting on a mind-boggling age of over 2000 million years. However, by the mid-20th century, most geologists were convinced of the validity of dating by radioactive decay. Radioactive decay gave the only precise method of obtaining the age of minerals, because it was the only method where the rates of production could be proved to be constant and where possible loss of products could be evaluated. There are three types of time relationships: age equivalence, relative age, and age in years. It is important to note, ﬁrst, that each of these relationships may be precise or imprecise, and second, that the various methods used to determine age may overlap. Some methods fall between these categories; for example, the 14 C method of radiometric dating gives ages in years, but these are not real calendar years due to variations in the initial carbon isotope ratios of the atmosphere through time, so the ages have to be calibrated against an absolute precise standard, such as tree-rings. Estimating the precision of a method also normally involves calibration with other methods. Calibration gives the ages in years of the geologic period and magnetic reversal timescales. Remember the difference between precision, accuracy, and resolution (Fig. 6.1).

7.1 Age equivalence Age equivalence was discussed in Chapter 6 on physical correlation. Precise age equivalence is possible due to isolated, very unusual events or processes that instantly deposit a unique sedimentary layer. If volcanic ash and meteorite impact layers can be proved to come from the same eruption or impact, then they give a precision of

correlation accurate to within days, even though the actual age is unknown. In theory, any unusual event can be used; for example, an unusually large storm might deposit a unique graded sandstone in a basin. However, in practice, it can be difﬁcult to prove uniqueness. Imprecise age equivalence is also given by the methods that give relative and absolute ages (see below). For example, fossil dating and magnetic reversals give a rough age equivalence without age, while radiometric dating gives ages in years, but with signiﬁcant experimental errors.

7.2 Relative ages Relative ages require a progressive change through time that can be measured. The precision of the dating depends on the measurable rate of change and its uniqueness. There are no really precise relative ages. Imprecise relative ages are given by physical relationships of strata, fossils, and a host of other methods including the changing chemistry and isotopic composition of the oceans and atmospheres through time, and migrating paleomagnetic poles and reversals. Precision varies both within and between methods. 7.2.1 Physical relationships Steno used cross-cutting relationships together with his three principles of superposition, initial horizontality, and initial continuity, to determine the stratigraphic succession in Tuscany (see Chapter 6). For example, for the rocks exposed in a quarry wall (Fig. 7.2a), superposition shows that shale 1, the sandstone, and shale 2 were laid down successively, and cross-cutting relationships show that the basalt dike intruded shale 1 and the sandstone before shale 2 was deposited (a rock is younger than the rocks it cuts, and older than the rocks that cut it). From these relationships we can work out a relative history: (i) deposition of shale 1, and then the sandstone; (ii) intrusion of the basalt dike; (iii) erosion (since the dike is an intrusive rock); and (iv) deposition of shale 2. Using the principles of initial horizontality and continuity we can trace these layers into nearby quarries (Fig. 7.2b,c), and amplify the stratigraphic succession by recording further layers (limestones) not found in the ﬁrst quarry. In deformed rocks, various criteria can show whether the rocks are the right way up or overturned. The actual time span of these events in indeterminable. Quite complicated histories can be worked out in

The time dimension 143 (a)

(b)

(c)

Limestone

Shale 2

Shale 2

Sst Sandstone

Shale 1

Shale 1

Basalt dike Fig. 7.2 Quarry exposures in Tuscany illustrating the use of Steno’s principles.

small areas using this procedure (it is recommended that you do one of the relative age exercises in a geology laboratory manual). However, it does not work if layer equivalence is in doubt, or if correlation is attempted outside individual sedimentary basins, or across major breaks. A succession of sandstone, shale, and limestone in Tuscany (central Italy) could be very different in age from a similar succession in Calabria (southern Italy) or in New York (USA).

7.2.2 Fossils (biostratigraphy) 7.2.2.1 Introduction The use of fossils was the ﬁrst successful method of determining the relative ages of distant rocks. In the late 18th century, fossils were used descriptively in the same way as rock types, to deﬁne and map a unit. While surveying canals in southwestern England, William Smith found that each rock unit contained characteristic and distinct fossils that always repeated in the same order in a vertical succession of strata. Using these fossils, Smith found that individual rock units could be placed in their correct stratigraphic position in scattered outcrops (Fig. 7.3). Thus began biostratigraphy, which is the correlation and dating of rocks with fossils. By

Sst

Limestone

1801, Smith had traced these layers across England and Wales to produce the ﬁrst comprehensive geologic map (Fuller 1969). Smith’s studies were entirely practical and he had no idea why the fossils changed upwards. Yet, within ﬁfty years of his discovery all the geologic systems had been named. These pre-dated Darwin’s Origin of Species (1859) and took no account of biological evolution. But they ﬁtted the 19th century view that God had created the world and populated it with organisms by special creation, though with multiple special creations (and extinctions). Since Darwin, most people attribute organic evolution to natural selection. Nevertheless, the geologic systems stand as proof that successive waves of extinction and evolution have punctuated the history of life. Good biostratigraphy requires three things: common fossils, good taxonomy, and the accurate location of these fossils in carefully measured sections. Common fossils are the luck of the draw, although persistent collecting can sometimes turn up enough of even rare taxa. Good taxonomy is more controllable and is essential in all biostratigraphic studies. This needs emphasizing. Good taxonomic studies of fossils are becoming rare, even though these are essential for most stratigraphy. Accurate location requires, ﬁrst, that the vertical changes in fossils be noted at one place and convenient

144 Chapter 7

Fig. 7.3 William Smith’s use of guide fossils and faunal succession to match beds: quarry, hill, and canal sections combine to give the section on the right (from Fenton & Fenton 1952, p. 74). (Copyright 1952 by C.L. Fenton and M.A. Fenton. Used by permission of Doubleday, a division of Random House, Inc.)

boundaries chosen, and second, that these changes and boundaries must be recognized at other places. The ideal end result is that the three-dimensional space occupied by the fossil species or assemblage can be mapped in the same way as the three-dimensional space of the rock units. The fossil space may correspond with the rock space if the organisms were controlled by the same factors controlling the deposition of the sediment, or the fossil space may appear completely independent of the rock units; furthermore, different taxa may give different fossil spaces (Fig. 7.4). The fossils showing these two extremes have been called FACIES FOSSILS and ZONE FOSSILS, and ascribed to environmental and evolutionary controls respectively, although environmental and evolutionary controls operate in both cases. Even the slowest evolving, envi-

ronmentally restricted organisms, such as Lingulid brachiopods, show some evolution, but this evolution is only sporadically recorded in the nearshore clastic sediments they lived in, and the time intervals capable of being discriminated are very long. Rapidly evolving, free-living, and widely distributed organisms, such as ammonites, allow ﬁner time divisions over wider areas. However, such organisms were still controlled by factors (temperature, salinity, oxygen content, etc.) in the environment they inhabited. Ammonites are not normally found in nearshore environments, are naturally absent from freshwater and non-marine environments, show faunal changes with depth and latitude, and differ (due to independent evolution) in isolated seas. Similarly, different trilobite “zones” are required for different parts of shelf seas, and without periodic intercalation of distinct

The time dimension 145

Fig. 7.4 Discordance of foraminifera zones with lithofacies, and discordance among some biostratigraphic horizons (from Krumbein & Sloss 1963, ﬁg. 10.9). (© 1951, 1963 by W.C. Krumbein and L.L. Sloss. Used with the permission of W.H. Freeman and Company.)

facies, it is difﬁcult to correlate the separate trilobite zones (Fig. 7.5 and Table 7.1). This problem is exacerbated when comparing zonal schemes erected on either side of major paleogeographic barriers (see Chapter 13). So, there is no sharp break between “facies” and “zone” fossils (Ludvigsen et al. 1986). Some fossils are more narrowly controlled by the local environment than others, and some are more rapidly evolving than others, but all are ecologically controlled within the span of their existence. 7.2.2.2 Fossil zones are based on the vertical ranges of individual species or assemblages of species in sections. The idea was based on bed-by-bed descriptions of the Jurassic of southern Germany by Quenstedt (1858). However, the concept was realized by Oppel (1856–8), one of Quenstedt’s pupils, who carefully noted the vertical ranges of all the fossils, particularly ammonites, in the sections. Species appeared and disappeared in an appar-

ZONES

ently random manner but distinctive faunal aggregates (zones) were the result of the overlapping ranges of the different ammonite species (Fig. 7.6). A zone is an assemblage of guide fossils, of which one is selected as the index species and names the zone (Arkell 1956). The base of each zone is deﬁned by the appearance of certain new species, and its top by the appearance of new species that deﬁne the base of the succeeding zone. So, only the zone bases are deﬁned; their tops correspond with the base of the next zone (Murphy 1977). Although Oppel (perhaps wisely) nowhere deﬁned a zone, it is perfectly clear that a zone refers to the vertical ranges of fossil species in rock sections, not to the strata themselves and not to abstract time intervals. In theory, a zone includes any bed deposited in any part of the world during the period in which its diagnostic fossils lived. This will vary for different assemblages based on different species. Exact correlations of zones based on different taxa can neither be expected, anticipated, nor reasonably proposed (Ludvigsen et al. 1986). Since

146 Chapter 7 Table 7.1 Trilobite zones for Ordovician lithofacies belts, northwest Canada (from Lenz et al. 1993, ﬁg. 19). (Reproduced will permission of the Geological Association of Canada.)

Oppel’s time, many different types of zones have been proposed: most are superﬂuous and can be eliminated with no great loss (Fig. 7.7). Assemblage zones and concurrent ranges zones are basically the same as Oppel zones, and all can simply be called “zones.” Abundance zones (or acme zones or epiboles) are strata that contain the greatest number of individuals of a particular species or group of species. They have no time value; they simply represent condensed horizons or population blooms, and may be at different stratigraphic levels in different places. Interval zones are deﬁned by two successive ﬁrst or last occurrences of unrelated taxa, and are poor relations of Oppel zones. Taxon range zones (or teilzones) are based on the ﬁrst and last occurrence of a single genus or species (or other taxonomic category) in individual sections, and are useless for comparing different sections. Lineage zones require usually futile attempts to recognize ancestor and descendent species, although they have been used to place “golden spikes” to deﬁne units; a conodont lineage deﬁnes the base of the Triassic System. The most useless concept in biostratigraphy is that of INDEX FOSSILS. These are abundant, distinctive species, widely distributed, and supposedly narrowly limited in

time. Index fossil horizons may mark widespread, shortlived events and are probably near-contemporary over small areas (although there is no way of proving this). They can be useful in tying nearby sections together (as William Smith did). However, index fossils are simply another lithological feature of the rock unit. The use of index fossils to deﬁne and date rock units should be abandoned. It discourages the essential reporting of speciﬁc locations and ranges of fossils in measured rock sections. 7.2.2.3 Good zone fossils The precision of fossil dating varies with type of organism, evolutionary rate, and the taxonomic level chosen. Any fossil group can be used for zoning. But since the aim is to recognize the smallest time intervals over the widest area, zone fossils should ideally have the following characteristics: 1 A relatively wide paleogeographic range, usually marking wide ecological tolerances. However, the present-day distribution of the zone fossils may be misleading if an originally restricted geographic distribution has been expanded by younger plate movements. Some

The time dimension 147

Fig. 7.5 Trilobite distributions in Middle Ordovician shelf deposits (from Lenz et al. 1993, ﬁg. 17). (Reproduced with permission of the Geological Association of Canada.)

Triassic Paciﬁc terranes evolved as oceanic plateau in the middle of the paleo-Paciﬁc Ocean, yet are now found in both Japan and North America. The present-day wide distribution of the Triassic faunas of these terranes is an artifact of plate movement. Likewise, the widely separated Lower Jurassic zones of Europe and eastern North America are similar because the Atlantic Ocean was then very narrow. 2 Limited vertical time range of species. Rapidly evolving lineages such as ammonites and graptolites are most suitable since the more rapid the evolution, the shorter in time are the zones – at least theoretically. Also preferable is an evolutionary lineage, so the ancestor and descendants of species can be plotted on a section; for example, the Ringsteadia–Pictonia–Rasenia–Aulacostephanus lineage of the Oxfordian–Kimmeridgian (Upper Jurassic) (Ziegler 1962). However, these are very

rare, and where present make the choice of zone boundaries more difﬁcult. Sharp breaks in species make the boundaries easier to deﬁne, but reduce the precision of correlation among sections. 3 Easily recognized changes of taxonomic features on the skeleton. The great variation in sutures, coiling, and ornamentation of ammonites allows their rapid evolution to be readily recognized and the species classiﬁed. The contemporary belemnites, on the other hand, require sectioning and further study before their equally rapid evolution can be worked out. So, belemnites are generally used as zone fossils only where ammonites are rare or absent. 4 Floating, swimming, or ﬂying forms. These forms are less likely to be controlled by speciﬁc bottom or surface conditions, and are more likely to be carried by waves, currents, or winds into graves in a variety of different

148 Chapter 7

Fig. 7.6 Ammonite distributions in the Sinemurian of southern Germany. Note that changes occur at small erosion or condensation horizons (modiﬁed from Bayer & McGhee 1985, ﬁg. 5). (Copyright (1985). Reproduced with permission of Springer-Verlag.)

environments. However, their distribution may be controlled by environmental differences in water masses and by where they get their food. Thus, no ammonite, graptolite, or conodont assemblage is ubiquitous, even on one ancient shelf. 5 Capable of being preserved. The commonest organism is of no use if it has no skeleton and cannot be fossilized except under unusual conditions. Thus, birds would be good for at least local zonation, but they are rarely preserved. 6 Relatively abundant. A rapidly evolving, easily classi-

ﬁed organism is of no use if it is too rare to plot ranges on rock sections. Thus, fossil ﬁsh might be useful if they were not so rare. 7.2.2.4 Extent of zones Given abundant and taxonomically diverse suitable fossils (such as those of the Lower Jurassic), detailed usable local zonal schemes can be erected with a high internal resolution. But how local is local? Some zonal schemes, such as the European Lower Jurassic ammonite zones

The time dimension 149 and the Ordovician–Silurian graptolite zones, seem recognizable over extensive areas. Other schemes, such as the conodont zonations of eastern North America, are restricted to smaller areas and have strange boundaries between faunal provinces. Why the differences and what are they due to? How can the distinct zonal schemes of different areas (based on the same organism or on different organisms) be reconciled and correlated?

Fundamentally, what is the precision of correlation between distinct zonal schemes (Fahraeus 1986)? Answers to this question involve understanding the biogeography of both modern and ancient organisms; how organisms spread over the earth and how separate faunal realms and provinces become established and disappear. I will defer considering these concepts until Part 3 of this book. The only comment I make here is that proposed uniform worldwide biostratigraphic classiﬁcations are misleading and inaccurate at best. All biostratigraphic units are regional to a greater or lesser extent. The classiﬁcations should reﬂect this or a spurious equivalence of rocks of different ages will result (Ludvigsen et al. 1986). 7.2.2.5 Graphic correlation with fossils

Fig. 7.7 Various types of zones (from Ludvigsen et al. 1986, ﬁg. 2). (© 1990 by W.H. Freeman and Company. Used with permission.)

In 1964, Alan Shaw proposed a graphic method of correlation whereby the ranges of fossils in two or more sections are correlated statistically. With two sections, the correlation can be done graphically. If two sections have the same ranges of fossil species, then a bivariate plot of the two sections gives a perfect 45° correlation line and the sections must have accumulated at the same rate (Fig. 7.8). Gentler or steeper correlation lines show that accumulation rates were constant but different in the two sections, and allow evaluation of relative rates of sedimentation for the sections (Fig. 7.9). A dog-leg indicates

Fig. 7.8 Bases and tops of ranges plot on a straight line at 45° C in two identical sections (from Prothero 1990, ﬁg. 10.12).

150 Chapter 7

Fig. 7.9 Range plots for two sections with constant but different sedimentation rates (from Carney & Pierce 1995, ﬁg. 1). (Reproduced with permission of SEPM.)

a change in relative accumulation rates (Fig. 7.10a). A ﬂat section indicates no accumulation (or erosion) in one section, and a diastem or unconformity (Fig. 7.10b). Real data, of course, rarely show such nice correlations; but range points that scatter off the line can be reexamined to see why. For example, a large scatter in the ranges of the same species in two sections of similar lithofacies requires explanation other than lithofacies control (Fig. 7.11). The scatter of the points about the lines shows how useful the fossils are for biostratigraphic discrimination. And the discrimination can be measured with linear regression and other statistical methods. Once satisﬁed that the taxa used give good results, the two best sections in an area are correlated to give a best reference section. Range extensions from overlapping sections can then be added to produce a composite standard section, which is a synthesis of many sections (Fig. 7.12). Individual ranges in other sections can then be compared with this composite standard to see if the section has any

breaks, or if the composite standard needs amending (Fig. 7.13). Shaw’s graphic correlation method has some great advantages over standard biostratigraphic methods (Carney & Pierce 1995). First, it forces people to carefully measure and describe sections and precisely locate fossils: if it did nothing else but foster these, it would be very worthwhile. Second, it provides easy visualization of the precision and accuracy of the taxa ranges used: all potential points of correlation can be seen and their usefulness evaluated. Third, relative rates of sedimentation can be inferred by comparing the section with a standard time-scale (Fig. 7.14). Fourth, the method can be extended to other methods of relative dating (see paleomagnetic reversals section). The method also has some problems if not used carefully. First, if two sections are similar lithologically, then spurious biofacies correlations may occur. This is one reason why simply plotting taxa ranges without including lithology will not do. Second, it can only be applied

The time dimension 151 (a)

(b)

Fig. 7.10 (a) Change in sedimentation rate and (b) diastem or unconformity in one section (from Doyle 1996, box 5.4). (© John Wiley & Sons Limited. Reproduced with permission.)

Fig. 7.11 Ranges of ammonite species in two limestone–shale sections in southern Germany and southern France, 600 km apart (from Ziegler 1962, ﬁg. 126).

152 Chapter 7

Fig. 7.12 Composite standard section (from Carney & Pierce 1995, ﬁg. 4). (Reproduced with permission of SEPM.) srs, standard reference section.

within one biogeographic region: separate standards are required for different biogeographic regions. Despite these caveats, graphic biostratigraphic correlation is one of the most practical ways of establishing a biostratigraphic standard in which the errors and limitations can be plainly seen. 7.2.3 Geochemistry

Fig. 7.13 Ranges in one section compared with a composite standard (from Sweet 1987, ﬁg. 3).

The use of chemical analyses to identify volcanic ash beds and correlate them has already been mentioned (p. 126). Chemical changes have also occurred in the atmosphere, oceans, and lakes through time. The mixing times for changes in the atmosphere and in interconnected oceans are of the order of a few thousand years at most, and are thus geologically instantaneous. Within one mixing body (including an individual lake), any chemical changes preservable in sediments will form a pattern through time that can be used for correlation if

The time dimension 153

Fig. 7.14 Correlation of ranges in a section with those of a standard time-scale (from Carney & Pierce 1995, ﬁg. 1). (Reproduced with permission of SEPM.)

the deposits are roughly the same age, and can be used for rough dating if the geochemical trend is in one direction through time. Magnesium, strontium, sodium, and other elements in water have been used in this way. However, geochemical changes are often slower than paleomagnetic reversals (see below), so the patterns are less distinct and more difﬁcult to identify and correlate accurately (Holland 1984). For example, the strontium (Sr) content of pelagic carbonates has varied throughout earth history (Fig. 7.15). Long-term variations are thought to be due to variations in the Sr/Ca ratio in sea water due to the relative amounts of calcite versus aragonite precipitated and also the intensity of submarine hydrothermal activity which supplies ﬂuids with low Sr/Ca ratios to the oceans. Globally, the relative increase in Sr/Ca ratios during the Lower Cretaceous and Miocene suggests a diminution in submarine hydrothermal activity at those

times. On this general trend are superimposed shorterterm ﬂuctuations, the main cause of which seems to be marine transgression and regression. However, the data are so scattered that only a very rough age estimate is possible on Sr/Ca ratios. 7.2.4 Stable isotopes Stable isotopes exist in well-deﬁned ratios in the oceans and atmosphere, controlled by dynamic processes. These ratios change as the importance of the various processes change through time. Patterns of changing isotope ratios can then be traced from one section to another in the same way as paleomagnetic reversals and geochemical changes. As with geochemical changes, atmosphere and ocean homogenization times determine the theoretical precision of the methods (Williams et al. 1988). Strontium isotope ratios vary over both

154 Chapter 7

Fig. 7.15 Sr/Ca ratio in unaltered Cretaceous–Tertiary planktonic foraminifera. The inset shows the mean Sr/Ca ratio at an 80% conﬁdence limit (from Holland 1984, ﬁg. 9.38). (Reproduced with permission of Princeton University Press.)

long and short time periods and can be plotted in the same way as Sr/Ca ratios (McArthur 1994). However, the best example is the use of oxygen isotopes to correlate the Quaternary ice ages. 7.2.4.1 Oxygen isotopes Oxygen has three stable isotopes: lighter and dominant 16 O (99.8%), heavier and rarer 18O (0.2%), and very rare 17O (which can be ignored). During evaporation, water (H2O) with the heavier oxygen isotope is preferentially left behind. So rainwater and freshwater are enriched in H216O. From this, two main correlation methods were developed using annual variations in ice layers and longer-term variations in sea water and the carbonate precipitated from it. First, evaporation and condensation depend on temperature, and ocean temperatures are far more stable than air temperatures. The isotopic composition of rainwater depends mostly on the temperature of evaporating sea water. So, precipitation in winter is isotopically

heavier than in summer. Ice cores from glaciers show isotopically light summer layers alternating with isotopically heavy winter layers; and, if undeformed, the pattern of layers can be used to date ice cores in the same way that annual rings can be used to date wood. Although mountain ice-caps only have young ice, Antarctica and Greenland have much older ice. The most spectacular result of oxygen isotope analysis of ice cores from Greenland was that the last interglacial showed extreme ﬂuctuations in d18O values, suggesting cold to warm alternations within decades, in contrast to the present interglacial. Second, there is normally a balance in the oceans between the loss of H216O by evaporation and the gain of H216O supplied by rivers; but if water is locked up in large continental ice sheets (or lakes), then ocean water becomes relatively enriched in H218O (Fig. 7.16). Carbonate deposited inorganically has an isotopic ratio of 18O to 16O close to that of the water from which it came (expressed as a deviation, d18O, from a standard). The ratio of 18O to 16O in organic carbonate may be close

The time dimension 155

Fig. 7.16 Oxygen isotopes in (a) non-glacial and (b) glacial times (from Bennett & Glasser 1996, p. 13). (1996 © John Wiley & Sons Limited. Reproduced with permission.)

to or signiﬁcantly different from this, depending on the physiology of the organism secreting the carbonate. Foraminifera secret stable calcite with an oxygen isotopic ratio close to that of the sea water in which they live. After studying several deep-sea cores, Emiliani (1955) suggested that the oxygen isotope variations he recorded reﬂected glacial–interglacial temperature changes. Later, Shackleton (1967) found that the isotopic composition was strongly controlled by the volume of ice locked up on land during glacial periods. The oxygen isotope variations are used to deﬁne oxygen isotope stages which reﬂect ice volumes and, by inference, warm and cold periods. Isotope stages are conventionally numbered backwards, with the interglacials odd-numbered (the present interglacial as 1) and the glacials even-numbered (the last glacial as 2). A detailed pattern of oxygen isotope changes now extends back to isotopic stage 102 at the Gauss–Matyuma paleomagnetic reversal boundary (at about 2.5 Ma), and ﬂuctuations extend back through the Tertiary (Fig. 7.17). The oceans are stratiﬁed. Near-surface sea water, down to a variable level (or thermocline), roughly follows seasonal changes of temperature. Planktonic foraminifera thus show combined temperature and icevolume effects. Deeper water is colder, and controlled by deep underﬂows of polar water. Benthonic foraminifera therefore have different (colder) isotopic ratios than planktonic foraminifera. The isotope differences developed during the Oligocene and mark the onset of extensive glacial conditions in Antarctica and the

development of a strongly thermally stratiﬁed ocean. Oxygen isotope shifts from benthic foraminifera are practically independent of temperature and reﬂect mostly ice-volume changes. Detailed benthic isotopic shifts can be closely matched from different seas and allow precise correlation of marine sediments (Fig. 7.18). 7.2.5 Paleomagnetism When magnetic materials cool below their Curie point (575 °C for magnetite, 200–400 °C for titanomagnetite) in the presence of a magnetic ﬁeld, they acquire a strong and stable remnant magnetization parallel to the earth’s magnetic ﬁeld (Butler 1992). This thermoremnant magnetization (TRM), if kept near earth’s surface temperatures, varies little over time. Sediments show directional remnant magnetization (DRM), due, mostly, to physical alignment of magnetic particles in the earth’s magnetic ﬁeld during deposition. This DRM is weak and difﬁcult to measure. All rocks are also gradually magnetized in the direction of the existing magnetic ﬁeld, acquiring a viscous remnant magnetization (VRM). In volcanic and other rocks with very strong TRM, the VRM acquired later is very much weaker and can be ignored; but in sediments, the DRM and VRM components may be of similar strength, and the measured remnant magnetization is a composite of primary DRM and secondary VRM. To isolate the primary DRM we have to remove the secondary

156 Chapter 7

Fig. 7.17 (a) Oxygen isotope stages back to about 2.5 Ma. Black bars are interglacials. (b) Composite marine d18O record for the Tertiary. Solid lines indicate ice sheets; dashed lines indicate early growth stages of ice sheets in east Antarctica (EA), west Antarctica (WA), and the northern hemisphere (NH) (from Wilson et al. 2000, ﬁgs 4.15, 4.16). (Reproduced with permission of Routledge.)

VRM with alternating current demagnetization methods. Remnant magnetism gives paleolatitudes of rock units, and also (due to reversals of the earth’s magnetic ﬁeld through time) provides a relative time-scale. Paleomagnetism has an advantage over biostratigraphy and geochemistry in that magnetic changes, as

far as we know, take place simultaneously over the entire world. 7.2.5.1 Paleomagnetic pole positions During the 1950s, people found that, although the earth’s magnetic pole had wandered across the earth

The time dimension 157

Fig. 7.18 Oxygen isotope variations in Cretaceous–Tertiary foraminifera from (a) tropical Central Paciﬁc and (b) temperate intermediate latitudes. Open triangles are planktonic foraminifera; solid triangles are benthonic foraminifera (from Brand & Morrison 1987, ﬁg. 14).

over time, polar wandering paths for different continents did not correspond (Fig. 7.19). This not only conﬁrmed continental drift, but gave a rough method of dating. If the PALEOMAGNETIC POLE of a sample of unknown age is plotted on a dated polar wandering curve for the same tectonic unit, its age can be obtained. The age is imprecise because pole positions have large errors (note the circles of conﬁdence in Fig. 7.19). 7.2.5.2 Magnetostratigraphy The geomagnetic ﬁeld has constantly reversed its polarity through time. These reversals form characteristic patterns backwards through time (each reversal is thought to take about 10,000 years). Remnant magnetization weakens during reversals, and may temporarily disappear – in which case charged particles would enter the upper atmosphere and greatly affect the earth’s environment.

Volcanic rocks give a strong and stable record for the reversals, and patterns can be determined from spreading ridges. The oceanic lithosphere takes on the earths’ magnetic ﬁeld in parallel strips as it forms at these ridges, and the patterns are easily measured outwards from the ridge crests and cross-correlated within the ocean basins (Fig. 7.20a,b). A composite pattern including all sea ﬂoors was completed in the 1970s and mapped in all oceans where strips could be detected. These patterns are calibrated to biostratigraphic ages of the overlying sediments and to radiometric dating of sea-ﬂoor samples or analogous lavas on land. The result is the standard reversal time-scale extending back to the oldest preserved oceanic lithosphere, which is of Middle Jurassic age (Fig. 7.20c). Sedimentary rocks give more continuous and detailed reversal records, making MAGNETOSTRATIGRAPHY one of the best ways of correlating thick, rapidly deposited and poorly fossiliferous sedimentary successions, which

158 Chapter 7

Fig. 7.19 Late Precambrian to Cambrian (C) apparent polar wandering paths for East Gondwana (circles and light shading; GF, Grenville orogen), and Laurentia (triangles and heavy lines) (from Powell et al. 1993, ﬁg. 1). (Copyright © 1993. Reproduced with permission of the publisher, the Geological Society of America, Boulder, Colorado, USA.)

are difﬁcult to correlate otherwise. Graphic correlation methods can be used to compare separate magnetic reversal sections in the same way as with fossils (Fig. 7.21). The main limitations of magnetic reversal methods for stratigraphy are as follows: 1 Increasing magnetic declination towards the poles makes the method increasingly inaccurate towards higher paleolatitudes, where measurement errors become very large. 2 The time patterns have initially to be recorded in sections, i.e. in space. Variations in sediment rate and/or time breaks may make it very difﬁcult to recognize particular patterns. 3 There is no way of independently ﬁnding which part of the time pattern is being determined. A rough idea of age (from fossils, radiometric dating, etc.) is necessary. 4 Many sediments are too coarse to allow consistent alignment of magnetized particles. Magnetic reversals are not usually determinable in coarse clastic rocks, for example.

5 Diagenesis and metamorphism may impose secondary magnetization. Despite these limitations, magnetostratigraphy is very important because it can be applied both to sediments with fossils and to igneous rocks datable radiometrically, and thus links biostratigraphy and radiometric dating.

7.3 Numerical methods (ages in years) Numerical methods require calculation from a known ﬁxed age, normally the present day. 7.3.1 Precise numerical ages Precise numerical ages are ages given in years, theoretically with no errors; and require counting constant time increments from a marker of precisely known age. This can be done in two ways. First, you can count backwards from the present, using a series of repetitive annual time markers such as tree rings. Second, you can count from

The time dimension 159 (a)

Fig. 7.20 (a) Pillow lavas forming, Hawaii (courtesy US Geological Survey); (b) magnetic anomaly pattern parallel to the Mid-Atlantic Ridge (from Strahler 1971, ﬁg. 25.28); (c) magnetic reversal time-scale, calibrated to dated fossil time-scale (from Ozima 1987, ﬁg. 4.4) (Copyright (1987). Reproduced with permission of Springer-Verlag.). Black indicates normal polarity; white indicates reversed polarity.

160 Chapter 7

Fig. 7.21 Magnetostratigraphy of an Indian section correlated with the magnetic reversal time-scale. Regression (r = 0.998) shows good correlation, no major breaks, and constant sedimentation (from Opdyke & Channell 1996, ﬁg. 5.9). (Copyright (1996), with permission from Elsevier.)

a deﬁned point, using recorded astronomical events. For example, the heliacal dawn rise of the star Sirius was recorded by Egyptian priests on the 16th day of the eighth month in the seventh year of the reign of the pharaoh Senwosret III. Since the heliacal rising moved through the Egyptian calendar in one Sothic cycle of 146 years, using a fairly complicated procedure it can be determined that the record refers to 1872 bc. Both of these techniques are subject to errors. 7.3.1.1 Tree rings Living trees form annual rings that show the tree’s age, and the TREE RINGS form patterns related to climatic variations during growth. By overlapping the patterns of living trees with those of dead trees, a dendrochronology can be established (Fig. 7.22). Trees that live a long time, such as the bristle-cone pine of the southwestern USA, can extend dendrochronology backwards over 7000 years. Although useful in itself, dendrochronology is most useful in calibrating radiocarbon ages to true calendar years (see below).

7.3.1.2 Varved clays Repeated couplets of coarser and ﬁner grained sediment accumulate annually in glacial lakes. True VARVED CLAYS form when spring melt agitates the water, causing ﬁne sand and silt to be transported into deeper water. This is followed by a summer phytoplankton bloom which dies in the autumn and (together with clay) settles to the bottom in the quiet water below the frozen winter surface. Like tree rings, the varves form a pattern which can be counted backwards through successively older lakes (Fig. 7.23). In Sweden, De Geer (1921) established a chronology back to 17,000 years BP, although more recent work has shown that the chronology is accurate only for the last 12,000 years. The main problem is establishing that rhythmites represent true varves and are not simply multiple summer inﬂow events. If enough organic matter is preserved in the varves they can be used to calibrate radiocarbon ages in the same way as tree rings, but further backwards in time. Alternatively, the age of an isolated varve sequence can be worked out

The time dimension 161

Fig. 7.22 Use of annual tree rings to establish a dendrochronology: on the left are overlapping records from living and dead trees; on the right, annual rings are deﬁned by cell-size variations in early and late wood each year.

(a)

Fig. 7.23 (a) Varved clays in a Quaternary section, and (b) a method of correlating varves.

with radiocarbon ages calibrated against tree rings. “Floating” varve sequences (i.e. those not tied to a marker date) can be used to correlate, but not date, older glacial lake deposits such as those in the Precambrian, Ordovician, and Permian.

Such simple counting methods are theoretically precise. However, in practice, all are subject to errors caused by mismatching separate events and sequences and, for dendrochronology and varve chronology, misidentifying annual layers.

162 Chapter 7 7.3.2 Imprecise numerical (radiometric) dating Imprecise numerical ages are given by the radiometric decay of unstable isotopes and by the effects of this decay on the lattice structure of minerals (ﬁssion tracks). Radioactive isotopes decay at a constant rate, regardless of the physical or chemical environment. In a closed system, such as a mineral, the decay products accumulate within the system. If the decay product can be identiﬁed and measured, simply dividing the accumulated amount by the rate of decay will give a numerical age – in theory! In 1905, Ernest Rutherford suggested that radioactive minerals could be used to date rocks. By 1907, Boltwood had calculated an astounding age of 2.2 billion years from a mineral from Sri Lanka. The radiometric decay formulas are very simple: dN dt = -lN

(7.1)

T12 = ln 2 l

(7.2)

where l is a decay constant; T1/2 is the half-life; N is the number of atoms of the radioactive element; and t is time. The precision, accuracy, ease, and speed of analysis has increased enormously since Boltwood’s time. Typical analytical errors are now less than 1% at the 95% conﬁdence (2f) level. However, although some methods give more precise ages than others (if you consider “precise” to mean within 1% of the actual age), there is still some statistical uncertainty in decay rates (now within 1%) and some analytical errors, so the dates obtained are imprecise to a greater or lesser extent (Renne et al. 2000). Good introductions to RADIOMETRIC DATING methods can be found in Dalrymple (1991), Dickin (1995), and Ozima (1987). The logarithmic decay of radioactive isotopes means that the amount of isotope lost during each successive time increment is half that of the increment before (Fig. 7.24). The HALF-LIFE gives an estimate of the usefulness of each isotope for dating. After six half-lives only 1.675% of the original isotope is left, and this decreases by only 0.8375% in the seventh increment. Due to this small change, analytical errors exceed increment changes and the dating errors become enormous. The half-life is thus a good measure of the limits of the various isotopic systems. Too short a half-life, and the method is limited to relatively recent times, and, unless the isotope is produced naturally in some way, it will long since have

Fig. 7.24 Logarithmic (isotopic) decay (from Dalrymple 1991, ﬁg. 3.2). (Reproduced with permission of Stanford University Press.)

disappeared. Too long a half-life, and there will be insufﬁcient decay to measure over geologically useful time increments. Because of this, relatively few naturally radioactive isotope decay paths are useful for geologic dating. All are based on constant isotope decay and measuring the amount of isotope left and/or decay products. It is difﬁcult to measure the actual value of isotopes, but using a mass spectrometer, it is comparatively easy to measure ratios accurately. So, ratios are used for calculations where possible. There were ﬁve assumptions made in the early days of radiometric dating, which were thought to be necessary for accurate dating. Some of these are essential; others are less so, and in fact can be used to infer ages of different processes operating on the same rock and/or mineral. 1 A constant decay rate of the parent isotope. This is a fundamental requirement that has now been proved by assumption 2. 2 The decay rate can be measured in the laboratory. Decay rates of the common isotopes are now known to within 2% or better for 87Rb and 147Sm, and within 1% for 40K, 235U, and 238U. All laboratories use the same decay constants and isotopic compositions, so that dates can be compared without tedious recalculations. 3 Parent isotopes and their decay products are retained in a closed system. This is a requirement for parent isotopes. It is rarely true for decay products. However, the differential loss of decay products for different isotopic systems at different temperatures means that the cooling history of a rock or mineral can be worked out (see below).

The time dimension 163 Table 7.2 Characteristics of main radiometric dating methods (from Nichols 1999, table 20.1). (Reproduced with permission of Blackwell Publishing Ltd.) Parent isotope

Daughter isotope

40

40

K

Ar

87Rb

87Sr

147Sm

143Nd

176Lu

176Hf

232

208

Th 235U 238U 14C

Pb

207Pb 206Pb 14N

Half-life (109 years) 1.250 48.8 1.06 3.5 14.01 0.704 4.468 5730 years

4 None of the decay product is in the system at the start. This can be overcome in some systems by using isotopic ratios and plotting several analyses in isochrons, for example in the rubidium–strontium and uranium– thorium–lead systems (see below). 5 The mineral or rock must be formed in relatively short time geologically, otherwise the age obtained will be an average of a long time period. For minerals, this is a necessary assumption; for rocks, the analysis of different minerals forming at different temperatures with different isotopic methods can give more details about its history, as in assumption 3. Of the 70 natural radioactive isotopes, 18 have long half-lives. The remaining 52 have short half-lives but are continually created by nuclear reactions in nature. For example, 14C (carbon) is continually created from 14N (nitrogen) by reaction with cosmic-ray neutrons. Here we will concentrate on the uses and problems of the commonly used uranium–lead (U-Pb), samarium– neodymium (Sm-Nd), rubidium–strontium (Rb-Sr), potassium–argon (K-Ar), and carbon-14 (14C) methods (Table 7.2). Other isotopic dating methods have the same problems. 7.3.2.1 Uranium–thorium–lead (U-Th-Pb) Uranium is a very rare element, so the U-Th-Pb dating method is restricted to minerals that have concentrated U, such as zircon. The method is based on the radioactive decay of three isotopes, 235U, 238U, and 232Th, which also decay to different isotopes of lead, respectively 207Pb, 206Pb, and 208Pb. Although the decay series between these end-members is complex, with many intermediate daughter products, in practice the decay series can be treated as if they were one-step decays. This is because each of the three decay series is independent of the

Parent’s decay constant (year-1) -10

4.692 ¥ 10 1.42 ¥ 10-11 6.54 ¥ 10-12 1.94 ¥ 10-11 4.95 ¥ 10-11 9.85 ¥ 10-10 1.55 ¥ 10-10 1.29 ¥ 10-4

Practical dating range (Ma) 1 to >4500 10 to >4500 >200 >200 10 to >4500 10 to >4500 10 to >4500 <80,000 years

others, and the half-lives of the intermediate daughters are insigniﬁcant compared to their parents. Three simple independent age calculations can be made from each of the three series. Since both decay and daughter products are solids, they tend to stay within the mineral after it crystallizes and give a date of crystallization. If these ages agree, then that is the age of crystallization of the mineral. More often than not, the three ages do not agree, because lead is a relatively volatile element and can be lost if the mineral is reheated later. Because of this problem, concordia diagrams are normally used to obtain ages. Concordia diagrams use the two isotopic ratio pairs 238 206 U/ Pb and 235U/207Pb. If the isotopic ratios of each of these pairs are plotted against one another for different ages, then the result is a curved line called a CONCORDIA PLOT, which is the locus of all concordant U-Pb ages that have not been disturbed since crystallization (Fig. 7.25a). Any undisturbed mineral will plot on this line, which is curved because the two uranium isotopes decay at different rates. If lead is lost after crystallization, then the amount of lead in the mineral will decrease, the Pb/U ratios will decrease, and the altered mineral grains will plot closer to the graph origin: how much closer depends on the lead loss. If all lead is lost then the Pb/U ratios become zero and the point will plot at the origin. The points will not lie on the concordia curve because lead loss does not fractionate the Pb isotopes. Lead loss is in the same isotopic proportion as existed when the mineral experienced loss of lead, and results in analyses of various grains plotting on a straight line connecting the concordia point with the origin. However, after the loss, lead will continue to be produced by radioactive decay and the straight line will move along the curve, with the lower intercept giving the time that the alteration took place (Fig. 7.25b). Thus, plotting the Pb/U ratios of

164 Chapter 7

Fig. 7.25 (a) U–Pb concordia curve with crystallization age (P) at time of lead loss; (b) discordia plot now, two billion years after lead loss (from Dalrymple 1991, ﬁg. 3.13). (Reproduced with permission of Stanford University Press.)

grains from an altered rock gives points on a straight line that intersects the concordia curve at two points: the ﬁrst gives the original crystallization date, while the second gives the alteration date. Kinks show two or more stages of alteration. 7.3.2.2 Samarium–neodymium (Sm-Nd) The decay of a rare earth element, such as U-Pb, involves two solid elements. So, if a large enough sample is analysed (maybe several tens of centimeters), it can treated as a closed system. 147Sm decays to 143Nd with a very long half-life of 1.06 ¥ 1011 (106,000 million years), which is even longer than for Rb-Sr decay. Some of the discrepancies in Sm-Nd, Rb-Sr, and U-Pb dates may in fact be due to uncertainties in the accepted Rb-Sr and Sm-Nd decay constants. The long half-life of 147Sm makes the method very useful for Precambrian rocks and also basic rocks which are difﬁcult to date by other methods as they lack U- and K-bearing minerals. 7.3.2.3 Rubidium–strontium (Rb-Sr) In Rb-Sr dating, 87Rb decays to 87Sr, and the method is easy to use, since Rb is a minor substitute for K in many

mineral structures. However, this method has a number of complications. First, the 87Rb to 87Sr decay has a very long half-life of 47,000 million years. This means that: (i) laboratory measurements of the decay constant are difﬁcult (the accepted value of the decay constant l, now 1.421 ¥ 1011 years, has varied over the years since the method was developed); (ii) young rocks and minerals (roughly less than 10 million years old) are difﬁcult to date accurately since the amount of 87Sr produced is so small; and (iii) errors are quite large for the same reason. Second, minerals contain inherited 87Sr (87Sri) as well as 87Sr produced by radioactive decay (87Sr) after the mineral crystallized. Thus Total 87Sr =

87

Sr +

87

Rb(elt - 1)

(7.3)

Rb is preferentially concentrated relative to Sr during both partial melting and fractional crystallization. So, rocks that have undergone differentiation, such as those in continental crust, have higher 87Sr/86Sri ratios than more “primitive” rocks, such as those from the mantle. Thus, the 87Sri value has to be known to calculate an age. Dividing equation (7.3) by the stable isotope 86Sr to express it in isotope ratios gives the following:

The time dimension 165

Fig. 7.26 (a) Rb–Sr isochron based on analysis of different materials after crystallization from a homogeneous magma; (b) wholerock isochron of three rocks crystallized from a differentiating magma. The intersect gives an initial Sr isotope ratio of 0.7075 (from Faure 1977, ﬁgs 6.1, 6.2). (This material is used by permission of John Wiley & Sons, Inc.)

87

Sr

86

Sr = ( 87 Sr

86

Sr)i + ( 87 Rb

86

Sr)(elt - 1)

(7.4)

This equation has two unknown quantities, the time the mineral formed, t, and the initial strontium isotope ratio, (87Sr/86Sr)i, when the mineral formed. In order to ﬁnd t, we must analyze the 87Sr/86Sr and 87Rb/86Sr ratios of at least two samples, insert the values in the formula, and solve the simultaneous equations to ﬁnd t and (87Sr/86Sr)i. If the minerals crystallized from a homogeneous magma, then the (87Sr/86Sr)i values for both minerals should be the same, and the amount of 87Sr produced should be proportional to the 87Rb in the mineral. Thus, plotting the measured 87Sr/86Sr ratios against the 87Rb/86Sr ratios should produce a straight line, whose x-intercept is the initial 87Sr/86Sr ratio and slope of the line is (et — 1) (Fig. 7.26a). All material of the same age and with a common (87Sr/86Sr)i should lie on this line. The ratios will lie off a straight line if something has disturbed the isotope system. So the narrowness of the scatter of values around the line is itself a good check on the reliability of the isochron, and can be evaluated statistically. Third, the 87Sr2+ ion is much larger than its parent 87Rb+ ion and has a double as opposed to a single change. 87Sr2+ produced by Rb decay is unstable within the Kbearing mineral lattices. It tends to leak out and can replace Ca2+ in adjacent minerals. So, by analyzing several whole-rock samples you can still get a good isochron,

even if the isotope system has been disturbed (Fig. 7.26b). You can go further. K-bearing minerals lose 87Sr2+ during metamorphism but retain 87Sr2+ after dropping below a blocking temperature. Plotting an isochron based on several analyses of one or more minerals gives the date the rock cooled below this temperature and the age of metamorphism, which may differ from the age of crystallization (Fig. 7.27). However, since the blocking temperatures of muscovite (500 ± 50 °C) and biotite (300 ± 50 °C) differ, cooling isochrons should really be based on one mineral (Dickin 1995). 7.3.2.4 Potassium–argon (K-Ar) In the K-Ar method, 40K decays to 40Ca and 40Ar. Most of the 40K (89.5%) decays to 40Ca; but this is the common isotope of calcium, so the radiogenic Ca is swamped by the non-radiogenic isotope and cannot be measured. Some 40K (10.5%) decays to 40Ar with a half-life of 1193 million years. This is long enough for the rocks that are as old as the earth to be dated, but short enough to allow measurable decay in rocks as young as ﬁve million years. Though 40K is only 0.012% of total potassium, potassium is a common element and so the method is applicable to many minerals and rocks. K-Ar is really the only good method of dating young basic rocks, and so it is the main method of calibrating the magnetic reversal time-scale.

166 Chapter 7

Fig. 7.27 (a) Whole-rock and (b) mineral isochrons for the Carn Chuinneag Complex, Scotland (from Faure 1977, ﬁg. 6.6). (This material is used by permission of John Wiley & Sons, Inc.)

Inherited 40Ar can be a problem, especially in young samples and samples with a low K content. However, the effect of this inherited 40Ar can be evaluated with the 40Ar-39Ar method, which is based on the production of 39 Ar from 39K by neutron irradiation. Dates are calculated from measured 40Ar/39Ar ratios after complete release of gas from irradiated samples. The main advantage of the method is that argon can be released by stepwise heating of irradiated samples, and a spectrum of dates can be calculated for each step. The shape of the curve enclosing the steps gives an estimate of the cooling rate of the sample and, in intrusive rocks, may help in understanding its unrooﬁng history (Fig. 7.28). Argon, an inert gas, is easily lost from the system at different temperatures for different minerals: about 700 °C for amphiboles and about 380 °C for biotite. These cooling ages (and those from Rb-Sr) from different minerals can help evaluate the history of a rock (Fig. 7.29). Rocks and minerals that lost some argon during thermal events may still release argon from retentive sites at high temperatures to give plateau dates around the time of crystallization. Rocks and minerals that have

excess 40Ar give anomalously old ages in the lowest temperature steps (hornblende in Fig. 7.29). 40 Ar/39Ar dating helps to evaluate the problems of inherited argon, determine whether a date obtained is valid, estimate the rate of cooling of a sample, and determine its thermal history. 7.3.2.5 Carbon-14 (14C) Carbon-14 is produced by cosmic bombardment of 14N in the upper atmosphere. The 14C reacts with oxygen to form 14CO2, which then mixes rapidly with isotopically stable carbon dioxide throughout the atmosphere. A dynamic ratio is set up between the rate of production of 14C and its decay back to 14N over time. Living organisms incorporate organic carbon in their tissues with the isotopic ratio of the contemporary atmosphere or water. After death, the 14C decays and the radioactivity of the dead organism declines. If the original amount of 14C in the living organism is known then the residual activity of the dead organism can be used to determine its age (Fig. 7.30).

The time dimension 167

Fig. 7.28 40Ar/39Ar age spectra of biotite sample, compared to predicted biotite cooling curves based on modeling argon diffusion in biotite (from Dickin 1995, ﬁg. 10.24). (Reproduced with permission of Cambridge University Press.)

Fig. 7.29 Age spectra for hornblende and biotite (modiﬁed from Dickin 1995, ﬁg. 10.26). (Reproduced with permission of Cambridge University Press.)

Fig. 7.30 14C dating based on residual activity; by measuring the remaining 14C activity (A), the 14C age can be read off the graph (from Faure 1977, ﬁg. 17.1). (This material is used by permission of John Wiley & Sons, Inc.)

168 Chapter 7

Fig. 7.31 Calibration of 14C dates with tree rings, 1100– 400 bc (from Bowman 1990, ﬁg. 23). (© Copyright The British Museum.)

The half-life of 14C is only 5730 ± 40 years, so the method is reasonably accurate only within the last 35,000 years or so (six half-lives). The method depends on knowing the original 14C/12C ratio in the living organism (since the decay product 14N is a common gas). Originally it was assumed that the 14C/12C ratio remained constant through time, but checking 14C dates against tree chronologies shows that this is not true. The 14C/12C ratio has changed markedly over the last 7000 years and a 14C age is thus not a calendar age. Uncalibrated 14C ages gave serious and systematic dating errors in archeology before the calibration problem was solved, and remain a problem prior to 7000 BP. Even within this period, initial 14C variations cause “ﬂat” sections where radiocarbon dates are indistinguishable over time spans as long as 400 years (Fig. 7.31). All materials used for 14C dating are easily contaminated by younger introduced material. Isotopic exchange with porewaters and intracellular growth of micro-organisms (such as fungal hyphae) are particularly difﬁcult to resolve. Radiocarbon dates are published with reference to a zero year (ad 1950) and using the original 5570 ± 30 years half-life; these dates can be converted to the new 5730 ± 40 years half-life by multiplying by 1.03.

Fig. 7.32 A cooling curve based on isotopic and ﬁssion track dating of different minerals (from Gallagher et al. 1998, ﬁg. 7). (With permission from the Annual Review of Earth and Planetary Sciences, vol. 26 © 1998 by Annual Reviews.)

7.3.2.6 Fission tracks A FISSION TRACK marks microscopic radiation damage caused by the breakdown of radioactive isotopes in a uranium-bearing mineral (e.g. zircon, sphene, apatite) or glass (e.g. obsidian) (Naeser & Naeser 1988). The rare isotope 238U has a ﬁssion half-life short enough to produce enough tracks to measure. Etching of polished sections enhances the tracks, which can then be counted microscopically. The age can be calculated from the amount of uranium and the number of spontaneous ﬁssion tracks in a mineral or glass. The relative abundance of 238U and 235U is constant in nature, so the easiest and most accurate way to determine the amount of uranium present is to create a new set of ﬁssion tracks by irradiating the sample with a known dose of neutrons. The track density due to induced ﬁssion of 235U (the commoner isotope) is related to the amount of uranium in the sample and the neutron dose it got in the reactor. Fission track dating gives cooling dates since tracks fade and disappear (anneal) on heating (the temperature varies with the individual mineral and its history, and is lower for glasses). Fission track ages, particularly in combination with other radiometric dates, are thus extremely useful in working out thermal histories

The time dimension 169

Fig. 7.33 Benthic foraminiferal change at the Cretaceous–Tertiary boundary, El Kef, Tunisia (from Speijer & Van Der Zwain 1996, ﬁg. 5).

Fig. 7.34 Iridum spike of impact clay in some sections and position on paleomagnetic time scale (from Archibald 1996, ﬁg. 7.1).

170 Chapter 7 (Fig. 7.32). By making assumptions about geothermal gradients and other factors, the erosional history of an area can be worked out. Fission track dates often have large errors, but they have no lower limit and can be used to roughly calibrate older radiocarbon dates.

7.4 Calibration of relative and numerical dates Relative time units based on fossils and other methods can be calibrated with numerical time units based on radiometric dating. The numerical ages of geologic systems are primarily based on dating igneous rocks associated with fossiliferous rocks. For example, the boundary between the Mesozoic and Cenozoic eras was deﬁned by the extinction of characteristic Mesozoic life forms such as dinosaurs and ammonites, and their eventual replacement by characteristic Cenozoic life forms such as mammals and teleost ﬁsh. Studying the fossil layers in more detail and correlating and calibrating the boundary can also help answer two questions. Did extinctions occur gradually or suddenly? Did land and sea extinctions occur at the same time?

concentration of certain platinum group elements (of which iridium was the easiest to measure) in the clay only by the impact of a large meteorite. Subsequently, the impact layer was found worldwide in many marine and land sections. If the impact origin was true, then the thin clay was a worldwide instantaneous time marker analogous to a worldwide volcanic ash – which some people thought it was! The precise impact clay marker allowed, for the ﬁrst time, detailed correlation of land and sea sections. If the impact devastated the earth, then extinctions should correlate with the impact clay. Both large fossils and more abundant microfossils were systematically collected from apparently continuous sections and related to the often thin impact clay. These fossils showed that although diversity dropped progressively in the later Cretaceous, the ﬁnal extinctions in both continental and marine sections occurred at the clay. 7.4.3 Imprecise numerical ages

For over two hundred years, stratigraphers and paleontologists have been reﬁning the relative dating of both land and sea sections with fossils. In many sections, Mesozoic forms disappear progressively towards the boundary, correlated with a drop in sea level. After a gap, new Cenozoic forms certainly come in progressively above the boundary. Restrictions of habitat and changing environments were thus popular explanations for the changeover. However, there is a preservation problem. There is usually an unconformity between the Cretaceous and Paleocene periods. Sections spanning the boundary are rare and simple collection area restriction is bound to result in fewer forms near the boundary. The boundary could not be precisely deﬁned in either land or sea sections, although detailed analysis showed that the boundary was sharp in continuous marine sections (Fig. 7.33).

Next, paleomagnetic reversal patterns in the sedimentary boundary sections could be related to the paleomagnetic reversal time-scale, which is based on reversals in basalt lavas: the impact layer lies within unit 29R (Fig. 7.34). The basalt lavas can be radiometrically dated on land using K-Ar dating. Although early results were unsatisfactory because of alteration and argon loss or gain, later results were more consistent and gave an excellent linear correlation between astrochronological and K-Ar and Ar-Ar ages for polarity transitions for the last four million years. This meant that the Ar-Ar ages on basalts as old as the Cretaceous were also reasonably accurate. Dating basalts within the 29R isochron gives an age of about 65 Ma. Impact glasses from the 180 km diameter impact crater at Chicxlub in Mexico gave Ar-Ar ages of 65.2 ± 0.4 Ma. So 65 million years is therefore the best rounded estimate of the numerical age of the Mesozoic–Cenozoic boundary. The extinctions on both land and sea were abrupt and synchronous and thus most plausibly caused by environmental changes associated with the large meteorite impact.

7.4.2 Relative precise correlation

7.4.4 Precise numerical age

In 1980, Luis Alvarez and colleagues published analyses of a thin clay layer at the marine Cretaceous–Cenozoic boundary in Italy. They could explain the peculiar

Unfortunately, there is no known method of obtaining a more precise numerical age for the Mesozoic–Cenozoic boundary than the 65.2 ± 0.4 Ma given above.

7.4.1 Relative imprecise dating

8 Basin analysis

8.1 8.2 8.3 8.4 8.5 8.6

Basin-ﬁll architecture Sediment provenance Paleocurrents and sediment dispersal Backstripping Paleothermometry Paleogeographic and paleotectonic maps

Strata accumulate in basins of various types, whose origin and evolution are controlled by tectonics, climate, sea level, and the organisms present; these, in turn, control sediment supply, accumulation, and geometry. Although tectonics is the fundamental control on basin formation and development, the type, nature, and geometry of the sediment ﬁll depend not only on tectonics but also on climate, sea-level changes, and the organisms present. For example, in a continental rift, a cold climate will give a different sedimentary ﬁlling to a hot climate, relative sea-level changes will cause facies migrations, and Precambrian sediments will differ from Cenozoic ones because of evolution. In ancient sedimentary basins, only the results of tectonics, climate, sea-level change and organisms can be studied. So, the type and extent of controlling factors can only be inferred from sediment character, distribution, thickness and geometry, together with fossils, tectonic structures, and other evidence. Ideally, these aspects should be analyzed before stratigraphic classiﬁ-

cation (Chapter 9), and before controls are evaluated (Chapters 10–13), difﬁcult as this is. This chapter summarizes the techniques of basin analysis, including basin-ﬁll architecture, sedimentary provenance, paleocurrents, backstripping, and paleothermometry (Allen & Allen 1990; Einsele 2000; Miall 2000).

8.1 Basin-ﬁll architecture Basin-ﬁll architecture is done by noting the threedimensional distribution of lithostratigraphic, facies, and time units, together with any major breaks between units formed under different environmental conditions. 8.1.1 Fence diagrams Fence diagrams are often based on boreholes, since outcrops are usually too few. The deﬁned units (such

172 Chapter 8

Fig. 8.1 Fence diagram of Cambrian sediments of the southwestern USA (from Baars 1983, p. 54).

as formations, lithofacies, or biofacies) are simply extrapolated across the area, noting changes and variations (Fig. 8.1). The diagrams are more useful if time-lines can be added to show facies changes. Twodimensional seismic sections can also be arranged as fence diagrams. 8.1.2 Sequence stratigraphy SEQUENCE STRATIGRAPHY uses the geometric relationships of strata on seismic sections to subdivide and correlate units (Fig. 8.2). Although initially based on twodimensional seismic sections, this technique is now used on three-dimensional seismic interpretations (see below) and on outcrops. A SEQUENCE lies between two unconformable surfaces (sequence boundaries) marked by progressive ONLAP of layers across a surface that frequently shows truncation of the underlying layers. Layers that merge with adjacent layers (usually because of condensation) show apparent truncation. Inclined layers resting progressively on a surface below (perhaps due to build-out of sediment)

show DOWNLAP. Layers that run across the top of these inclined surfaces show TOPLAP. On an ideal seismic section, two very different sediment packages can be identiﬁed bounded by basal onlap and basal downlap. These simple descriptive terms are very useful in delineating and mapping seismic packages across a basin. However, since the seismic reﬂection lines often cut across facies distributions, they have been interpreted as time-lines as well as reﬂector lines (Fig. 8.3). On this basis, the descriptive packages can be interpreted as genetic depositional intervals bounded by time-lines. More controversially, they are used to deﬁne a stratigraphic system of depositional sequences and systems tracts controlled by worldwide (eustatic) sea-level changes (see Chapter 9). 8.1.3 Three-dimensional seismic reconstructions Three-dimensional reconstructions from twodimensional seismic sections can now easily be done

Basin analysis 173

Fig. 8.2 Terminology of geometric relationships in sequence stratigraphy. Discontinuity terms are in bold (from Vail 1987, ﬁg. 1).

Fig. 8.3 Relationships of seismic reﬂector lines to facies, and four variables that affect basin ﬁlls: tectonics (subsidence), eustasy, climate, and sediment supply (from Vail 1987, ﬁg. 3).

using computers, with emphasis placed on important layers or units (Plate 8.1). Furthermore, other information (e.g. oil and gas character, porosity, etc., from wells) can also be added and individual surfaces analyzed and

interpreted (Plate 8.2) (for examples, see Weimer & Davis 1996). In virtual-reality laboratories, three-dimensional displays can be manipulated to investigate the basin from

174 Chapter 8

Fig. 8.4 (a) Combined facies, thickness, and paleocurrents, Cretaceous, Mississippi Basin (from Pryor 1960, ﬁg. 18). (b) Isopach maps of successive units, showing syndepositional fault control, Tertiary, Baise Basin, China (from Changsong et al. 1991, ﬁg. 4). Isopachs in meters. (Reproduced with permission of Blackwell Publishing Ltd.)

all aspects (Hart 1999). However, as convincing as such displays are (and they are often used to convince doubting executives about the hydrocarbon potential of a basin), they are not reality but interpretations based on limited data (Plate 8.3). Furthermore, they are based on data that have already been interpreted to give the seismic sections or boreholes (see Chapter 6). It is impossible to illustrate three-dimensional seismic reconstructions adequately in a text: examples can be seen on the internet; or better, a visit to a virtual-reality laboratory should be arranged to get the full ﬂavour of the amazing possibilities of the technique. Successive threedimensional analyses restored through time are called four-dimensional analyses. These can be animated and manipulated for changes in any factor or group of factors, from oil/gas contacts to the evolution of

an entire basin. The examples here have been limited to hydrocarbon-bearing basins, but the principles apply to every aspect of stratigraphy, from changes in river channels to the evolution of orogens (Plate 8.4). 8.1.4 Isopach maps ISOPACH maps show thickness variations. They can be compiled for an entire basin ﬁll, and include other information such as facies distributions and paleocurrents (Fig. 8.4a). However, they are more informative if they are done for successive depositional units, in which case they can show how changing distributions of sediment relate to tectonics or other factors (Fig. 8.4b).

Basin analysis 175

Fig. 8.4 Continued.

8.2 Sediment provenance Sediment PROVENANCE can be inferred by comparing the composition of the sediment with possible sources. This not only helps in determining the source rocks but also how they were weathered and transported. For example, the similar proportions of local unstable source rocks and angular pebbles in alluvial fan breccias within small Permian rift basins suggests rapid physical weathering, transportation, and deposition (Fig. 8.5).

8.3 Paleocurrents and sediment dispersal PALEOCURRENT

analysis shows where sediments were

transported from and to, and thus how the basin was ﬁlled in (Fig. 8.6a). Sediment dispersal may be evaluated from the relative sizes of clasts (Fig. 8.6b), although including paleocurrent and facies analysis should give better results (Fig. 8.4).

8.4 Backstripping BACKSTRIPPING is a technique for analyzing the subsidence and sedimentary history of a basin. It is a quantitative way of working out where and when sediment accumulated, and whether the sediment supply kept pace, or not, with subsidence and sea-level changes. Such studies are now routinely used to infer which controls dominated basin development.

176 Chapter 8

Fig. 8.5 Composition variation in Permian breccias, southern Scotland, and inferred feeder valleys. Clast percentage symbols as for bedrock (from Brookﬁeld 1978, ﬁg. 4). (Copyright (1978), with permission from Elsevier Science.)

Basin analysis 177

Fig. 8.6 Paleocurrents and sediment dispersal: (a) paleocurrent roses for Devonian river sandstones, Somerset Island, Canada; spread of directions indicates meandering streams (from Miall 1990, ﬁg. 2.19) (Copyright (1990). Reproduced with permission of Springer-Verlag.); (b) sediment dispersal from pebble size, Carboniferous conglomerates, Pennsylvania (from Meckel 1967, ﬁg. 5). (Reproduced with permission of the Geological Society.)

The main factors controlling post-rifting subsidence of both intracratonic and passive margin basins are thermal contraction and sedimentary loading (McKenzie 1978). Factors such as compaction, paleobathymetry, and long-term changes in sea level all contribute to the subsidence, but their combined effects are small compared with thermal contraction and sediment loading (Watts 1982). The preservation of thick marine shelf sequences ultimately depends almost entirely on the thermal subsidence of thinned continental lithosphere. As subsidence rates asymptotically decay and accommodation space is reduced, shelf sediments can build up to near sea level. The time constant from rifting

to negligible subsidence is about 150 Ma (McKenzie 1978). This is independent of spreading rates, which produce either a narrow or a wide shelf. Net subsidence thus places a deﬁnite limit on the total vertical thickness of sediment that can accumulate in any one place. For greatly thinned continental crust (thinning factors of 10), about 9 km of clastic sediment can accumulate at progressively slowing rates. For thicker or longer sequences to form, rifting must be repeated in the same or nearby areas. If the basin is not ﬁlled by 150 Ma after the start of rifting, then inﬁlling will proceed laterally until it is building out successive overlapping sediment prisms. Each prism is capped by thin

178 Chapter 8

Fig. 8.7 Thermal subsidence control of large-scale stratal geometry and preservation (from Brookﬁeld 1994, ﬁg. 2). (Copyright (1994), with permission from Elsevier Science.)

sedimentary successions, broken by unconformities (Fig. 8.7). The simplest backstripping uses one vertical section (a well log is ideal) to determine the tectonic displacement of the crust in a given time period. All sediment above the topmost horizon is removed and the underlying layers decompacted. The height difference of the column after removal of this layer gives the vertical displacement of the crust since the horizon was deposited. Most of this is isostatic subsidence due to the weight of the sediment (Chapter 10), some may reﬂect eustatic changes in sea level (Chapter 11), and some is due to sediment compaction (Fig. 8.8). Balancing the pressure at a reference horizon below loaded and unloaded mantle–crust columns gives the following relationships for equilibrium:

pw.g.Wd + ps.g.S * + pc.g.T = Y . pw.g + pc.g.T + x. pm.g

(8.1)

where Wd is water depth; S* is decompacted sediment thickness; Y is tectonic subsidence; g is gravity; T is mean crustal thickness; and Pc, Pw, and Ps are the densities of crust, water, and decompacted sediment, respectively. The equation can also be written as follows: x = Wd + S * + T - (Y + ds + T )

(8.2)

where ds is the sea-level change. From this, the tectonic subsidence or backstripping equation can be derived:

Y = Wd + S * [( pm - ps ) ( pm - pw )] - ds[ pm ( pm - pw )]

(8.3)

(Its isostatic basis is noted in more detail in Chapter 10.) The ﬁrst term in the equation is a water-depth term, the second is a sediment-loading term, and the third is a sealevel loading term.

Basin analysis 179

Fig. 8.9 Time slices in the subsidence history of a small area. Water depth, sedimentation rate, and compaction vary with time, and each sediment increment is compacted beneath succeeding increments, which require removal during backstripping (from Miall 1990, ﬁg. 2.18). (Copyright (1990). Reproduced with permission of Springer-Verlag.)

Fig. 8.8 Basic relationship between a loaded and an unloaded (backstripped) section (from Miall 1990, ﬁg. 2.19). (Copyright (1990). Reproduced with permission of Springer-Verlag.)

However, S*, the decompacted sediment thickness, has to be estimated at each stage from the decompaction equation: S * = S (1 - q) (1 - q *)

(8.4)

where S is the thickness of the compacted sediment, q is the porosity of the compacted sediment, and q* is the porosity of the uncompacted sediment. The decompacted thickness can be calculated from a standard porosity versus depth curve. For example, a 100 m thick shale has porosities of 0.12 at a depth of 3 km and 0.70 at the surface. The decompacted thickness is 100 (1 - 0.12)/1 - 0.70) or 293 m, i.e. more than 2.5 times the compacted thickness. Other

lithologies, such as sandstones and limestones, undergo much less compaction. Backstripping of successive layers from the top to the bottom of a section gives various time slices showing the tectonic subsidence history (Fig. 8.9). The mid-points of these time slices can be plotted to give tectonic subsidence curves for different sections of wells within a basin, or even of different structural units (Fig. 8.10). Many such subsidence curves assume constant paleobathymetry (often approximating zero for shallow shelf deposits) and no compaction (which is reasonable since most compaction occurs within a few hundred meters of the surface). If curves deviate from logarithmic subsidence, then reduced sedimentation rates, increasing paleobathymetry, signiﬁcant compaction, erosion, or combinations of these, have to be invoked. Nothing is ever simple in geology. For example, in the Mesozoic St George’s Channel Basin of the Irish Sea, two wells with intra-Triassic horizons both show anomalies in their subsidence curves (Fig. 8.11). In the basin center, under the assumption of constant paleobathymetry, there seems to be slight basement uplift in borehole “A” during the Middle Triassic (Carnian); while to the south, there is over 100 m of apparent basement uplift in borehole “B” during the Lower Triassic (Scythian). This is unlikely as the whole area was undergoing extension at the time and there is no evidence of even local compression. The anomalies can be explained in two ways: by subsidence without sedimentation, or by initial basement height above sea level. The subsidence curves for the later Triassic are consistent with constant sediment surface elevation.

180 Chapter 8

Fig. 8.10 Tectonic subsidence curves for structural units of the Himalayas. Note the steep curve for southern Tibet from late Permian times which marks the splitting of northern Tibet from Gondwanaland and the formation of the Neotethys ocean; all other curves are on continental, or thinned continental, crust (from Brookﬁeld 1993, ﬁg. 20). (Copyright (1993), with permission from Elsevier Science.)

The most likely explanation (shown on Fig. 8.11 as corrected for paleobathymetry) is that the basement was originally around 100 m in the center and 400 m at the margins. Attempts are sometimes made to infer paleodepths from worldwide eustatic curves, but the futility of this is described in Chapters 10 and 11. The latter part of this paragraph, in fact, anticipates interpretations of basin evolution requiring the concepts of Chapters 10–13. Subsidence curves can be compared for a series of sections or wells and integrated to show the threedimensional subsidence history of a large area (Fig. 8.12). Subsidence histories through time can also

be reconstructed through backstripping, and related to regional or local causes, such as faults (Plate 8.5). However, all such models depend on the accuracy and precision of, and assumptions underlying, the basic data. Although they look very nice, they are still interpretations of limited data. As with all such modeling, you should never forget the programmers’ warning: garbage in, garbage out.

8.5 Paleothermometry PALEOTHERMOMETRY

is an attempt to reconstruct the

Basin analysis 181

SE

SW

Fig. 8.11 Map of St George’s Channel basin, Irish Sea, UK, and tectonic subsidence curves for boreholes A and B, under constant paleobathymetry and corrected for paleobathymetry (courtesy of M.J. Welch, School of Earth Sciences, Birmingham University).

thermal distributions during basin ﬁlling, and evaluate how hydrocarbons formed. The oil-generation geothermal window is limited, and if the rock temperatures rose above this for any length of time, all the oil will be converted to gas (Fig. 8.13). Certain techniques, such as ﬁssion track dating, can partially reconstruct the thermal history of rocks and see if they stayed at temperatures low enough to host oil.

Fission tracks anneal (disappear) at temperatures that vary with the mineral. A maximum temperature for a particular time period can be determined, and sometimes tested with other methods, such as vitrinite reﬂectance (or increased “coaliﬁcation” of organic matter with temperature) or conodont colours (conodonts darken with increasing temperature). For example, the thermal history of a well in the Irish Sea indicates that it

182 Chapter 8

Fig. 8.12 Changing subsidence history of the St George’s Channel basin during the Triassic (courtesy of M.J. Welch, School of Earth Sciences, Birmingham University).

Fig. 8.13 Hydrocarbon yield for increasing temperature and standard depth equivalents (from Hunt 1996, ﬁg. 5.14).

Fig. 8.14 Thermal history of Permo-Carboniferous sediments in borehole 110/7-2, Irish Sea, UK: (a) location of borehole; (b) cross-section showing borehole penetrating horst; (c) ﬁssion track ages at base of borehole versus stratigraphic ages; (d) paleotemperatures from annealing temperatures of apatite ﬁssion track ages (AFTA) and vitrinite reﬂectance (VR); (e) inferred thermal history of Permo-Carboniferous strata. Samples cooled from a paleotemperature high of about 110 °C in the early Tertiary (from Green et al., 1997, ﬁg. 2). (Reproduced with permission of the Geological Society.)

184 Chapter 8

Fig. 8.15 (a) Devonian clastic thickness and relation to Appalachia (from Kay & Colbert 1965, ﬁg. 11.10) (Copyright © 1965. This material is used by permission of John Wiley & Sons, Inc.); (b) Baker’s (1936) map showing the ﬁt of Transatlantic Paleozoic orogenies with continental drift (from Du Toit 1937, ﬁg. 3).

reached temperatures above 100 °C about 60 Ma, in the early Tertiary (Fig. 8.14). The high temperatures may be the result of deep burial followed by exhumation, or due to the early Tertiary phase of heating associated with the Iceland hotspot and the splitting of Greenland from northwestern Europe. In any case, almost all the Permian–Triassic rocks of the Irish Sea at some time went through the oil window, and gas is the only hydrocarbon commonly found there.

8.6 Paleogeographic and paleotectonic maps Paleogeographic and paleotectonic maps are attempts to show the evolution of sedimentary basins through time. They incorporate all available information and can be done on a local scale, or on a continent-wide scale. Like all such maps, they should be viewed with great suspicion unless all the assumptions on which they are

based are stated and carefully evaluated. For example, Schuchert’s (1955) classic paleogeographic maps of North America assume ﬁxed continents. To get around the problem of an eastern source for coarsegrained Paleozoic clastics in the Appalachians (only the deep Atlantic Ocean lies to the east), Schuchert simply put a land mass on his maps, called Appalachia, where he knew one had to be. The idea was immensely inﬂuential: Appalachia made it into all historical geology textbooks (Fig. 8.15a) – despite the isostatic impossibility of the idea, despite Schuchert being very familiar with Wegener’s continental drift theory, despite the fact that all major objections to continental drift had already been disposed of by the 1930s (Oreskes 1999), and despite the fact that continental drift maps that explained the data better had already been published (Fig. 8.15b). Judge theories on their ability to explain facts, not on others’ opinions about their validity.

Basin analysis 185

Fig. 8.16 Paleotectonic–paleogeographic maps of the North Atlantic, Carboniferous to Upper Permian (redrawn from Ziegler 1987, plates 6–8).

Paleotectonic maps now usually incorporate plate reconstructions on a large scale. Like Schuchert’s (1955) maps, they often include many unstated assumptions. For example, Ziegler’s (1987) paleogeo-

graphic–paleotectonic maps of the North Atlantic region include inferred major transcurrent faulting based on paleomagnetically determined major plate motions (Fig. 8.16).

9 Stratigraphic systems

9.1 9.2 9.3 9.4

Development of the stratigraphic system Cycle stratigraphy Genetic sequence stratigraphy The current system

“The tendency has always been strong to believe that whatever received a name must be an entity of being, having an independent existence of its own. And if no real entity answering to the name could be found, men did not for that reason suppose that none existed, but imagined that it was something abstruse and mysterious.” John Stuart Mill (1843) A System of Logic

appreciate or be able to criticize the current systems. The second section outlines the concepts of rhythmic or cyclical sedimentation. The third section outlines genetic sequence stratigraphy, while the fourth section notes some current systems and their good and bad points.

9.1 Development of the stratigraphic system The main aim of stratigraphy is the reconstruction of historical change which, on earth, is the result of the interaction of cyclical processes (e.g. the surface and plate tectonic cycles) with unidirectional changes (e.g. loss of heat and biological change). The challenge of stratigraphic systems is to incorporate these conﬂicting “time’s cycle” and “time’s arrow” views (Gould 1987) into a coherent classiﬁcation. This chapter considers the various ways in which stratigraphic systems were erected in the past, and how this led to the systems now used. The ﬁrst section summarizes the history of stratigraphic systems because, unless you know what has been tried and tested in the past, you are unlikely to

The problems of dividing local sections into manageable rock and time units (Chapter 5) are magniﬁed when studying extensive areas and long time periods outside one sedimentary basin. Some consistent philosophy or system is required. Although there is no need to proliferate names on that account, some hierarchical grouping of categories is needed, analogous to biological nomenclature. But what should be the basis of such nomenclature and of stratigraphic systems? During the late 17th and early 18th century, local rock successions were established around mining areas. In 1725, Strachey published a geologic cross-section of

Stratigraphic systems 187 Table 9.1 Werner’s classiﬁcation of rocks in Germany.

Alluvial Volcanic Floetz Transitional Primitive

Lithology

Fossils

Sand, gravel, loom, peat, bog deposits Lava, pumice, ash, tuff Stratiﬁed units of limestone, sandstone, Clay, shale, coal, rock salt, gypsum Chemical precipitates interbedded with Clastic sedimentary rocks Granite, porphyry, gneiss, basalt, serpentine marble, quartzite

Scarce to abundant; many reworked from Floetz

successive tilted and faulted coal layers in the Somerset coalﬁeld in the same way as we do today. Extending observations over more extensive areas meant that larger units were needed. People recognized that the older rocks generally formed the cores of mountains, with successively younger rocks at lower elevations. In 1756, after extensive and thorough ﬁeldwork in southern Germany, Lehman divided rocks into three groups. High mountains, such as the Alps, were made of highly deformed strata, without fossils, but with a lot of useful minerals, which he called Gang-Geburge (gang means “ore”). Below these, were lower mountains and hills with more gently inclined, fossiliferous strata, which he called Flotz-Geburge. According to Lehman, these formed when Noah’s Flood swept sediment and organisms off the Gang-Geburge to accumulate on the mountain ﬂanks, where they hardened and became rock. The even lower and softer sands and clays of the north German hills and plain formed by natural accidents, such as volcanic eruptions, earthquakes, landslides, storms, etc. Slightly later, Arduini in Italy named these divisions Primary, Secondary and Tertiary. These observations led to Werner’s synthesis (published in 1787) in which a Universal Ocean gradually receded, depositing successive layers with differing compositions over time. He inserted a Transition unit between Lehman’s Gangand Flotz-Geburge to describe highly deformed, though slightly fossiliferous, rocks. Thus, at the end of the 18th century, strata were classiﬁed according to their composition, which was supposed to reﬂect successive deposition in a Universal Ocean. The rock classiﬁcation was also a relative time classiﬁcation, and the two need not be separated (Table 9.1). At the same time, French biologists were ﬁnding out that Werner’s rock units had very different fossils in them (or no fossils in the Primary). It was believed that successive catastrophes had killed off earlier life forms, and the bare earth was then repopulated by newly created life forms (Fig. 9.1).

Generally abundant Generally abundant Lacking

The schemes of both Werner and Cuvier could be reconciled with the Christian religion. In fact, the geologic systems we now use were primarily erected in the early 19th century using these ideas. D’Orbigny invented the biostratigraphic stage to overcome the lithostratigraphic confusion of the French Jurassic. And Oppel’s zones gave a way of organizing the Jurassic strata of the whole of Western Europe in time, independent of the proliferating local rock names (Bork 1984). Even Hutton’s unconformities could be ﬁtted into the scheme, marking the convulsive rise and fall of sea level during successive great ﬂoods that deﬁned the geologic systems. By this time, geologists such as Elie de Beaumont had proposed that these convulsions raised mountains during the main changes in life forms. A grand synthesis involving great catastrophes separating the quieter geologic systems and eras seemed appropriate and appeared to ﬁt the rock record (Grabau 1913). The rise and fall of sea level and the deformation of mountains were considered to be worldwide events that allowed the stratigraphic column to be divided up, especially after acceptance of the glacial theory and the low but indisputable age for the earth calculated by physicists (Fig. 9.2). At the same time as systems based on worldwide periodic convulsions were being developed, the counter-argument of gradual changes punctuated by great time gaps was being developed by Hutton, Lyell, and others. Lyell divided up the Cenozoic based on the relative percentage of living taxa in the fossils from successively older strata. The obvious breaks in the percentages meant that the Cenozoic units were separated by long time gaps proportional to the differences (Fig. 9.3). Meanwhile, Gressly and others were also undermining the Universal Ocean by pointing out that rock units of different character and with different fossils graded into one another and were contemporaneous, rather than successive layers through time.

Fig. 9.1 The youngest catastrophe – the Deluge by Francis Danby (1840). (© Tate, London 2003.)

Fig. 9.2 Worldwide sea-level and mountain-building phases (from Grabau 1936).

Stratigraphic systems 189

Fig. 9.3 Lyell’s Tertiary chronology based on a uniform rate of biological change, with enormous inferred time gaps between preserved units (from Rudwick 1976, ﬁg. 4.5A).

Uniformitarians were unhappy with the short tens of millions of years for the age of the earth calculated by the physicist Lord Kelvin, but there was little they could do about it except grumble. The discovery of radioactivity in 1895 led to accurate numerical dating of rocks, and pushed the age of the earth back by more than two orders of magnitude to 4.6 billion years. Despite this, people who accepted uniformitarian ideas somehow saw no problem in accepting universal oceans with distinct faunas, and periodic worldwide mountain building. Dana (1870) proposed that mountain chains formed during successive epochs of folding, uplift, and metamorphism, separated by long intervals of accumulation of crustal tension. Ulrich (1911) saw the Paleozoic of eastern North America as a succession of distinct sedimentary lenses with distinct faunas deposited in isolated, ever-changing marine basins (Byers 2001). It is difﬁcult to believe now that the concepts of both facies and biostratigraphy were rejected for almost the entire ﬁrst half of the 20th century in North

America (Johnson 1985; Berry 1987). Even as late as the mid-20th century, stratigraphic systems were being erected based on the Universal Ocean model of sedimentary and biologic change (Kay 1960), or on worldwide orogenic pulses corresponding to sea-level changes (Stille 1950). This persists to this day in some variants of sequence stratigraphy (see Section 9.3). However, by the mid-20th century a reaction to these worldwide schemes set in, and many people denied regularity to the tectonic, eustatic, climatic, and biologic processes that control the stratigraphic record (e.g. Gilluly 1949). Nevertheless, regular repetitions of sediments were used to infer cyclicity and cyclical controlling factors, and schemes based on these were erected throughout the 19th and 20th centuries. Such schemes can be regular or irregular, intrinsic to the system (autocyclic) or imposed from outside (allocyclic). They have a long history, are still a part of stratigraphy, and require a separate section.

190 Chapter 9 9.2 Cycle stratigraphy “The ﬂight of time is measured by the weaving of composite rhythms . . . the stratigraphic series constitutes a record, written in tablets of stone, of the lesser and greater waves of change which have pulsed through geologic time.” (Joseph Barrell (1917, p. 746)) Accepting that strata tend to come in packages separated by disconformities, there are three unresolved questions that have been debated since the days of Werner and Hutton in the 18th century, through Gressly, Barrell, Grabau, and Ulrich in the late 19th and early 20th centuries, and right up to the present controversies of sequence stratigraphy (Weller 1964; Miall & Miall 2000). First, what type of repetitive packages exist? Second, are these controlled by normal and local changes within a basin (autocyclic processes) or by basin-wide or even worldwide changes (allocyclic processes)? Third, how regular are the processes and what controls their periodicity? All these questions are bound up with the recognition and correlation of cycles at different scales (Duff et al. 1967). 9.2.1 Cycle classiﬁcations A CYCLE is an ordered succession of rock types that repeat in a predictable pattern, though the pattern has to be neither complete nor fully developed. Short-term, simple cycles are the annual varves in seasonally frozen lakes, and the repetitive ﬁning-upwards channel–overbank cycles of meandering streams which may take several years to several hundred years to form (Fig. 5.8). The cycles are often incomplete due to non-deposition, erosion, or incomplete migration. Similar cycles are the graded beds deposited by recurrent turbidity currents and storms, and the carbonate–clay couplets of some successions. A CYCLOTHEM is a more complicated cycle, and the term was speciﬁcally coined for Pennsylvanian sediments of North America. A whole plethora of terms such as megacyclothems, hypercyclothems, magnacycles, and other hierarchical schemes such as cycles, mesocycles, macrocycles, and megacycles and so on, have been used from time to time. None with success, however, because the hierarchy cannot be related to magnitudes and scales of controlling factors. The same problem applies to the ﬁrst-, second-, third-order cycles of sequence stratigraphy (Carter et al. 1991). The longest cycles are the tens of millions of years long unconformitybounded sequences of cratons, attributed to tectonics

or eustasy (Sloss 1963), and the even longer climatic icehouse–greenhouse cycles (Fischer 1986). These sequence cycles are deﬁned on the repetitive sequences of beds. Other cycles include “thickness cycles” in which beds repeat at equal vertical intervals, and “time cycles” in which beds repeat at equal time intervals (Schwarzacher 1975). With the exception of Milankovitch cycles which are a sort of time cycle, neither thickness nor time cycles are now much used because they seem to have no possible basis in reality. Few geologists can think of processes that give equal space or time increments. Sequential cycles can be studied statistically, for example with Markov chain analysis. Although popular at one time, such analyses do not seem to help understanding much and are now little used (Miall 1997). 9.2.2 Composite cycles occur where cycles of different wavelengths are superimposed upon one another (Barrell 1917). Good examples of these are the various Milankovitch climatic cycles superimposed on one another to give a very complex stratigraphic sequence in even relatively complete Quaternary deposits (Imbrie & Imbrie 1979). The problems with cycles are, ﬁrst, recognizing them; second, correlating them from one area to another; third, identifying what caused them; and fourth, convincing others that they are not just ﬁgments of your own imagination – which is where statistical analysis helps (Schwarzacher 1975). To illustrate the problems, we will look at two examples: limestone–marl cycles, and cyclothems.

COMPOSITE CYCLES

9.2.2.1 Limestone–marl cycles Many interbedded limestones and shales show rhythmic oscillations in carbonate content, resulting in carbonate-rich and carbonate-poor beds. Some of the most thoroughly studied examples come from the classic Lower Jurassic of Dorset, England, where there are repeating sequences of calcareous mudstones, bituminous shales, and clayey limestones (Fig. 9.4) (Hallam 1964). Various processes contribute to cycle development. Some are caused by changes in carbonate and/or clay supply as proved by changing trace-fossil assemblages (Fig. 9.4c), but others show diagenetic segregation (Fig. 9.4b). Are the cycles due to tectonic, eustatic, climatic,

Stratigraphic systems 191 (b)

(a)

(c)

Fig. 9.4 Limestone–shale cycles of the Lower Jurassic of Dorset: (a) cliff section; (b) diagenetic segregations; (c) typical Blue Lias cycle (courtesy I.M. West, School of Ocean and Earth Sciences, Southampton University, UK).

192 Chapter 9 or biological rhythms (Chapters 10–13)? Climatic and associated biologic changes are more likely causes because the cycles are small-scale and repeat many times (House 1985). High-resolution time series derived from magnetic susceptibility measurements were attributed to Milankovitch astronomical periodicities (Chapter 12), which, in turn, were used to calibrate the Jurassic time-scale (Weedon et al. 1999). However, some individual cycles change laterally, some beds are impersistent, and there is some tectonic control (Hallam 1986; Wignall 2001). 9.2.2.2 Cyclothems Cyclothems were originally described from the Pennsylvanian and early Permian of the central United States. Each has a characteristic vertical scheme of rock units that record marine invasion and retreat. A typical Pennsylvanian cyclothem is up to 20 m thick and is repeated many times. It starts with a disconformity, which is overlain by a basal sandstone passing up into underclay, coal, black shale, marine limestone, gray marine shale, algal limestone and shale, and brackish shale (Fig. 9.5). The cyclothem concept is one of recurrent patterns of sedimentation separated by breaks, but few cyclothems contain all the units in any one place (Wanless & Weller 1932). Some Pennsylvanian cyclothems can deﬁnitely be correlated over extensive regions and are attributed to allocyclic causes, such as eustatic sea-level changes triggered by the advance and retreat of southern hemisphere ice sheets (Ramsbottom 1979). However, the ﬁlling of basins is usually episodic in time and space at several orders of magnitude, and many cyclothems are simply local autocyclic facies changes (George 1978). Switching delta channels can form overlapping depositional lenses on a small scale within one delta lobe; and switching delta lobes may form intricate three-dimensional disconformity surfaces between successive sedimentary packages, as in the Cenozoic Mississippi delta (Fig. 6.5). Autocyclic processes are often difﬁcult to separate from allocyclic controls (Yang et al. 1998). And, of course, these depend on what scale you are considering: an allocycle for the earth is an autocycle for the rest of the solar system. Allocyclic processes such as relative changes of sea level may have ultimate controls varying from climatic to eustatic to tectonic (Fig. 9.6). Like Lias cycles, Pennsylvanian cyclothems have been related to sea-level changes forced by Milankovitch

variations in the earth’s orbit. Such explanations for cyclicity are increasingly used in stratigraphy (Hinnov 2000), to the extent that a hierarchy is used to classify cycles, based on their duration, and their possible causes inferred (Table 9.2). However, without good time correlation and dating it is impossible to decide what controlled cyclothem development (Klein 1993). All such correlations ultimately depend on accurate numerical dating of the strata involved, which is usually too imprecise to decide on the periodicity interval. For example, the Pennsylvanian is often taken as 34 million years long, but it may be only 19 million years long. With this substantial 44% decrease, the Milankovitch correlations fall apart and other causes for the cyclicity must be invoked (Klein 1990).

9.3 Genetic sequence stratigraphy Genetic sequence stratigraphy is an outgrowth of the use of seismic proﬁling and sequence stratigraphy in basin analysis (Chapter 8). The fundamental assumptions of genetic sequence stratigraphy are that seismic reﬂectors mark time-lines or isochronous surfaces and that large-scale facies geometry is primarily caused by sea-level changes (Mitchum et al. 1977). Genetic sequence stratigraphy has two parts. The ﬁrst is concerned with the generation of sequences and systems tracts by changes in relative sea level. The second attempts to correlate these sea-level changes globally assuming they are eustatic. 9.3.1 Generation of sequences A sort of genetic sequence stratigraphy has been used for many years (Sloss 1988), but it is popular now because of its usefulness in basin analysis (Vail 1987; Galloway 1989; Christie-Blick & Driscoll 1995). A geometric terminology is used to describe stratal relationships based on seismic sections (Fig. 8.2), while a genetic terminology derived from this interprets the sedimentary bodies deﬁned (Van Wagoner et al. 1987) (Fig. 9.7). A DEPOSITIONAL SEQUENCE is a relatively conformable succession of genetically related strata bounded by unconformities and their correlative conformities. The bounding unconformities mark drops in base levels of erosion caused by relative changes of sea level; while the depositional sequence forms in the ACCOMMODATION SPACE

Stratigraphic systems 193

Fig. 9.5 (a) Kansas-type cyclothem (from Moore 1964, ﬁg. 1) (Reproduced with permission of the Kansas Geological Survey.); (b) Permo-Carboniferous cyclothems in San Juan River, Colorado; (c) coal (25 cm) between underclay and shale, Pennsylvania.

194 Chapter 9 (b)

(c)

Fig. 9.5 Continued.

Fig. 9.6 Three major types of North American cyclothem and their controls (from Miall 1997, ﬁg. 8.30). (Copyright (1997). Reproduced with permission of Springer-Verlag.)

Table 9.2 Stratigraphic cycles and their probable (possible) causes (from Prothero 1990, table 8.1). (© 1990 by W.H. Freeman and Company. Used with permission.) Type (Vail et al. 1977)

Other terms

Duration (million years)

Probable cause

First order

Supercycles (Fischer 1981)

200–400

Second order

Sequence (Sloss 1963) Synthem (Ramsbottom 1979) Mesothem (Ramsbottom 1979)

10–100

Cyclothem (Wanless & Weller 1932)

0.2–0.5

Major eustatic cycles caused by formation and breakup of supercontinents Eustatic cycles induced by volume changes in global mid-oceanic spreading ridge system Possibly produced by ridge changes and/or continental ice growth and decay Rapid eustatic ﬂuctuations induced by growth and decay of continental ice sheets, growth and abandonment of deltas

Third order

Fourth order

1–10

Fig. 9.7 Genetic sequence stratigraphy: (a) units deﬁned in space; (b) units deﬁned in time; (c) interaction of eustasy and tectonics to give relative sea-level curve (from Prothero & Schwab 1996, ﬁg. 17.19 and Vail 1987, ﬁg. 7, rearranged).

196 Chapter 9

Fig. 9.7 Continued.

created by rising base levels (see Fig. 5.11). The time relationship diagram (Fig. 9.7b) is very useful in emphasizing stratal breaks. A TYPE 1 SEQUENCE BOUNDARY has subaerial erosion with channels incised into the shelf and deposition of sediment at the base of the slope. It is caused by the base level dropping below the shelf edge. A TYPE 2 SEQUENCE BOUNDARY lacks extensive subaerial erosion and valley incision since the base level does not drop below the shelf edge (Fig. 9.7). These deﬁne type 1 and type 2 sequences. Depositional sequences are further divided into SYSTEMS TRACTS, based on mutual relationships, inferred geometry and facies, and inferred environment. A LOWSTAND SYSTEMS TRACT forms at the base of a slope above a type 1 sequence boundary and often consists of lowstand fans and turbidites. A TRANSGRESSIVE SYSTEMS TRACT forms as sea level rises, and has backstepping shallowing-upwards cycles showing an overall transgression. INCISED VALLEY FILLS also form as the base level rises, and shallowing-upwards cycles tend to be progressively thinner because of sediment starvation. This culminates in a MAXIMUM FLOODING SURFACE of condensed facies, which separates the transgressive systems tract from the overlying regressive highstand systems tracts. A HIGHSTAND SYSTEMS TRACT has upwardly and outwardly building sigmoidal strata passing up into packets of erosion-bounded subaerial strata. It forms as accommodation space is ﬁlled in during stagnant or slowly falling relative sea level. A SHELF-MARGIN SYSTEMS TRACT

may form as a prograding wedge below the shelf edge. Smaller shallowing-upwards cycles within the sequences are known as PARASEQUENCES and PARASEQUENCE SETS. These are bounded by marine ﬂooding surfaces and are often claimed to be controlled by smaller relative sealevel changes (Fig. 9.8). Such assumptions underlie the system of punctuated aggradational cycles (PACs) (Goodwin & Anderson 1985). Both sequences and parasequences are often incomplete vertically, and especially laterally where the interaction of tectonics, eustasy, climate, and sediment supply may generate completely different and incompatible sequences even within the same basin. For example, while one delta lobe is building out a highstand systems tract, an adjacent lobe may be subsiding and forming a transgressive systems tract (Fig. 6.4). Grounded glaciers entering a lake reduce accommodation space to zero beneath the ice, forming sequence boundaries, while at the same time causing lake levels to rise and forming transgressive system tracts in open water (Brookﬁeld & Martini 1999). Two distinct and out-of-phase sequence stratigraphies have to be erected within the same sedimentary basin. The same problems occur in basins formed by ﬂexural loading (see Chapter 10). 9.3.2 Global sea level There is an unfortunate tendency to emphasize either

Stratigraphic systems 197

Fig. 9.8 Typical shallowing-upward cycles (parasequences and parasequence sets) in a transgressive systems tract, Lower Cretaceous (Aptian), UK (from Ruffell & Wach 1998, ﬁg. 9). (Reproduced with permission of Blackwell Publishing Ltd.)

local sedimentary and tectonic control or ubiquitous eustatic control, without realizing that all operate together at different scales (see Chapters 10–13). Thus, major ﬂooding surfaces are attributed to eustatic rises of sea level and used to construct global eustatic charts of unbelievable precision which are claimed to be useful in intercontinental correlation (Fig. 9.9). However, “cratons, their margins, and their interior basins do not just lie there passively waiting to be encroached upon by rising sea levels or laid bare to erosion as sea level falls” (Sloss 1988). Tectonic deformation occurs at many scales, and can be orders of magnitude larger and faster than nonglacial eustasy (Dewey & Pitman 1998). It is important to recognize the scale and extent of bounding surfaces and how they formed since sequence stratigraphy is based on genetic depositional packages, not on lithological units as in conventional stratigraphy (Miall 1996). Widespread sequence boundaries caused by worldwide eustatic sea-level changes of the same magnitude are conceptually impossible (Chapter 11) and contradicted by what data exist (Miall 1986). On three widely separated Jurassic–

Cretaceous passive margins where biostratigraphic correlation is good, only two out of 27 sequence boundaries are demonstrably synchronous: global synchroneity of all but two Mesozoic sequence boundaries is thus an illusion (Hubbard 1988). The late Wisconsin– Holocene Mississippi River deposited a type 1 sequence that includes lowstand, transgressive, and highstand system tracts within the last 30,000 years. Such time spans are far beyond the hundreds of thousands to millions of years’ resolution of pre-Quaternary stratigraphy (Boyd 1989). Also, tectonic ﬂexure on one continental margin distributes accommodation space according to lithospheric rigidity. Higher rigidities produce wide shelves which favour type 1 sequences; while lower rigidities produce narrow shelves which favour type 2 sequences (Reynolds et al. 1991) (Fig. 9.6).

9.4 The current system “The need to keep time and rock distinct in our thoughts is

198 Chapter 9

Fig. 9.9 Part of global eustatic correlation chart for the Zuni (Cretaceous) supersequence (from Haq et al. 1988). (Reproduced with permission of SEPM.) obvious, and to the extent that this elaborate terminology has led to clariﬁcation of thought it has served a useful purpose. But beyond that it is unnecessary. No one uses it, nor ever will.” (W.J. Arkell (1956))

All stratigraphic systems are attempts to order and clarify rock and time relationships among strata. By the mid- to late 20th century, various stratigraphic guides and codes had been published to regulate the naming

Stratigraphic systems 199 Table 9.3 Categories and ranks of stratigraphic units (North American Commission on Stratigraphic Nomenclature 1983, table 2). Material units Lithodemic

Supergroup Group

Supersulte Suite

Formation

Lithodeme

Member (or lens or tongue) Bed(s) or ﬂow(s)

Magnetopolarity Complex

Lithostratigraphic

Polarity Superzone Polarity zone

Polarity Subzone

Biostratigraphic

Pedostratigraphic

Allostratigraphic Allogroup

Biozone (interval assemblage or abundance) Subbiozone

Geosol

Alloformation

Allomember

Temporal and related chronostratigraphic units Geochronologic, geochronometric

Polarity chronostratigraphic

Polarity chronologic

Eonothem Erathem (Supersystem) System (Subsystem) Series Stage (Substage) Chronozone

Eon Era (Superperiod) Period (Subperiod) Epoch Age (Subage) Chron

Polarity Superchronozone

Polarity Superchron

Polarity Chronozone

Polarity Chron

Polarity Subchronozone

Polarity Subchron

Diachronic Episode

Diachron

Chronostratigraphic

Phase Span Cline

Fundamental units are italicized.

of rock and time units, and to establish systems of naming that would be internally consistent. These vary from relatively simple to perhaps excessively complicated. There are two approaches to stratigraphic nomenclature (as expressed in various guides and codes). The ﬁrst provides a minimum of rules for general application: “it is an unhelpful exercise to set up rules so unyielding that signiﬁcant numbers of stratigraphers will simply go their own way without following them” (Holland et al. 1978: this guide is 18 pages long). The second attempts to prescribe intricate rules for all rocks and occasions: “The North American Stratigraphic Code seeks to describe explicit practices for classifying and naming all formally deﬁned geologic units” (North American Commission on Stratigraphic Nomenclature 1983: this code is 35 pages long). Note also the ominous difference between the terms “Guide” and “Code” used in each (see

Ager 1984). It should be pretty obvious by now which approach I favor; and this chapter follows neither the North American Code (1983) (Table 9.3) nor the International Guide (Salvador 1994) (Table 9.4) in its categories. The North American code separates material units from temporal and related chronostratigraphic units; but then confuses matters by putting relative time units (polarity and fossil zones) in the material category and mixing chronostratigraphic and geochronologic categories. The International Code has separate biostratigraphic, magnetostratigraphic, and chronostratigraphic categories, though the ﬁrst two are simply methods of obtaining the latter. Read Harland (1992) for further pertinent criticism. Here, I use rock, relative time, and numerical time divisions. The basic concepts and terminology required for understanding stratigraphic systems were introduced in Chapters 5–7. A basic division into rock

200 Chapter 9 Table 9.4 Categories and unit terms of stratigraphy (modiﬁed from Murphy & Salvador 1999, table 1). (Reproduced with permission of the IUGS.) Stratigraphic categories

Principal stratigraphic unit terms

Lithostratigraphic

Group Formation Member Bed(s), Flow(s)

Unconformity-bounded

Synthem

Biostratigraphic

Biozones: Range zones Interval zones Lineage zones Assemblage zones Abundance zones Other kinds of biozones

Magnetostratigraphic polarity Other (informal) stratigraphic categories (mineralogic, stable isotope, environmental, seismic, etc.)

Polarity zone -zone (with appropriate preﬁx)

(lithostratigraphic), relative time (chronostratigraphic), and numerical time (geochronologic) units is quite adequate for all purposes (Table 9.5). Chronostratigraphic units are erected on the time of formation of rock bodies and, as currently used and named, are based on relative time methods. These are then calibrated with numerical (geochronologic) methods, though in the Precambrian, where relative methods are very coarse, units are deﬁned numerically and are purely geochronologic (see Section 9.4.3). Systems of stratigraphic nomenclature need to be ﬂexible and adjustable, not rigid and ﬁxed; and if they confuse simple relationships, they should be discarded. There is no formal stratigraphic code in this book, because different places use different codes. Furthermore, such stratigraphic codes can be mesmerizing and destructive of thought: witness the unfortunate practice of erecting large numbers of stratigraphic terms and then arguing about their precise boundaries and lateral extent, without either measuring detailed sections or considering the paleoenvironments, geometry, and methods of correlating and dating of the strata (Kay 1937). 9.4.1 Rock units (lithostratigraphy) Most large rock exposures been subdivided into a hierarchy of manageable units in order to describe them suc-

Chronostratigraphic Eonothem Erathem System Series Stage Substage (Chronozone)

Equivalent geochronologic units

Eon Era Period Epoch Age Subage (or Age) (Chron)

cinctly. It does not matter whether the rocks are igneous, metamorphic, or sedimentary; they all have spatial dimensions that are not necessarily of the same age everywhere. Rock units are described from either surface or subsurface sections: surface sections are preferable because they are, in principle, easier to revisit and resample. Initially, subdivisions are made according to macroscopic ﬁeld observations such as rock type and structure, color, fossil content, and so on, although these subdivisions may change after more detailed laboratory and other analyses. Tabular units can be classiﬁed and grouped in hierarchies with the standard bed–member (and variants)–formation–group (and supergroup, etc.) terms of Chapter 5 (lithostratigraphy). It is unnecessary to separate out “ﬂow” for a lava bed. Massive and irregular bodies, such as many igneous intrusions and metamorphic rocks and salt, can be referred to as complexes and their geometry noted with suitable terms such as salt domes (in the North American guide they are separated as lithodemic units). The formation is the basic lithostratigraphic unit. Formations can be combined into groups, which are useful for smaller-scale mapping. Members are lithologically separable subdivisions of a formation, which may be tabular or form lenses (enclosed within other lithologies) or tongues (extending as a wedge from the main body of the forma-

Stratigraphic systems 201 Table 9.5 Units of rock, relative time, and numerical time. Rock units (lithostratigraphy)

Relative time units (chronostratigraphy)

Numerical time units (geochronology)

Supergroup Group Formation Member, etc. Bed, etc.

Eonothem (e.g. Phanerozoic) Erathem (e.g. Paleozoic) System (e.g. Ordovician) Series (e.g. Viruan) Stage (e.g. Uhaku) Zone (e.g. G. teretiusculus)

Eon Era Period Epoch Age Chron

tion). A bed is a distinctive layer that is often deﬁned because it is useful in local or regional correlation. A complex has diverse rock types, often metamorphic or igneous. Dunbar and Rogers (1957, pp. 257–66) nicely outlined the complexities of deﬁning lithostratigraphic units. Other, perhaps unnecessary, categories in the current guides and codes are lithodemic units, unconformitybounded units, allostratigraphic units, and pedostratigraphic units (Harland 1992). Lithodemic units do not show superposition and are intrusive and metamorphic bodies as noted above. Unconformity-bounded units (International Guide) are the same as allostratigraphic units (North American Code): both are bodies of rock and are adequately covered by standard lithostratigraphic terms such as group. Pedostratigraphic units are alteration horizons, and naming them as a separate stratigraphic category opens various cans of worms for other surfaces. In general, there is all-round agreement on the deﬁnition of rock units, if not on the terminology required. This is not the case for relative time units. 9.4.2 Relative time units (chronostratigraphy) Chronostratigraphy gives older/younger relationships without numerical dates (Chapter 7). Chronostratigraphic units are bodies of rock formed during a deﬁned interval of time and are all relative time units (International Guide 1999). Boundaries of chronostratigraphic units should be deﬁned on the basis of observable paleontological, physical, or chemical features of the rocks (North American Code). It is convenient to use the terms zone, stage, series, and system for chronostratigraphic divisions since these terms were the ones ﬁrst used to deﬁne relative time units based on fossils. Such terms can be extended to polarity reversals and other

methods with modifying preﬁxes to show what method is being used. High resolution may be possible with event layers, such as impact layers, ash falls, major storms, and so on, that are traceable regionally or worldwide, and useful chronostratigraphic units can be deﬁned between them (Kauffman 1988). Separating individual events can be a problem, although impact and volcanic ash layers do have distinguishable individual mineralogical and geochemical properties (Huff et al. 1996). The end-Cretaceous impact layer is recognizable worldwide, and ash layers allow chronostratigraphic divisions of about 40,000 years in the Cretaceous of western North America. 9.4.2.1 Biostratigraphy Biostratigraphy is now based on the evolution and extinction of organisms through time, and biostratigraphic units are deﬁned on fossils. The zone is the basic biostratigraphic unit, corresponding to the range of one or more taxa. Zonal schemes based on individual taxa cannot be useful across many biofacies, and zones based on different taxa need not correspond. Many schemes ﬁt fairly well simply because many organisms disappear at roughly synchronous extinction horizons, ranges are extended to existing boundaries, or boundaries of different zonal schemes are arbitrarily correlated (Fig. 9.10). Although zones are the only biostratigraphic units in the current North American and International Codes, fossils are still used as the standard for deﬁning most stage, series, and system boundaries, and these are used as relative time units here. The bases of units are deﬁned on placing a “golden spike” at the ﬁrst occurrence of a taxa in a Global Stratotype Section and Point (GSSP) (Cowie 1986; see Aubrey et al. 1999, for various problems with this). The boundaries of all such units must be imprecise and variable, depending on the environmental limits and vagaries of preservation of the taxa used.

202 Chapter 9

Fig. 9.10 Different zonal schemes for the Middle Jurassic of southeastern France (from Cotillon 1992, ﬁg. 18). (Copyright (1992). Reproduced with permission of Springer-Verlag.)

For example, the base of the Triassic is deﬁned as the ﬁrst appearance datum (FAD) of the conodont species Hindeodus parvus (Kozur et Pjatakova) at the Meishan section in China (Yin Hongfu et al. 2001). Other fossils would give a different position. Nevertheless, carbon and oxygen isotope changes allow pattern correlation with other sections, and volcanic ashes can be dated radiometrically. The system boundary (biostratigraphic) can be correlated with geochemical patterns (chemostratigraphic), and the period boundary dated (numerical), though with disputable conﬁdence limits. However, the basal Triassic system boundary, identiﬁed on different fossil taxa, magnetic reversals, geochemistry, and isotopes, cannot exactly correspond anywhere in the world. As Aristotle wrote (4th century bc), “it is the mark of an instructed mind to rest satisﬁed with the precision which the nature of the subject admits, and not to seek exactness where only an approximation of the truth is possible.”

9.4.2.2 Magnetostratigraphy Magnetostratigraphy is based on the patterns generated by reversals in the earth’s magnetic ﬁeld, which take around 10,000 years to complete (Chapter 7). Such reversals are inherently more precise than biostratigraphy because they are worldwide and synchronous. The main difﬁculties are that the patterns can repeat (and a rough initial dating with other methods is necessary), and the analyses are expensive. The basic unit is the magnetostratigraphic polarity zone, which can be subdivided into subzones, and grouped as superzones. The magnetostratigraphic system is a bit haphazard. Quaternary to Cretaceous reversals are given numbers (though the Quaternary has names for zones and subzones) but Early Cretaceous to Jurassic reversals were independently given M0, M1, etc., designation after the long, Late Cretaceous, quiet normal period (Fig. 7.20c).

Stratigraphic systems 203 Ma 60

Van Hinte (1976)

Odin et al. (1982)

M

M

Harland et al. (1982)

Kent and Gradstein (1985)

Ma 66.5

M

70

Maastrichtian 74.5

Ca Ca 80

Sa Co

90

T

Ca

Campanian

Sa

Sa Co T

Co T

Santonian Coniacian Turonian

Ce

Ce

Cenomanian

Ce 100

97.5 Al

Al

110

Al

Albian

Ap

Aptian

Ap Ap

113

Ba Ha

Ba

119

120 Ha

130

84 87.5 88.5 91

V Be

V

Ba

Be

Ha

Barremian 124 Hauterivian 131

V

Valanginian 138

140 Fig. 9.11 Different radiometric scales for the Cretaceous (from Cotillon 1992, ﬁg. 15). (Copyright (1992). Reproduced with permission of SpringerVerlag.)

Be

Berriasian 144

150

9.4.2.3 Chemostratigraphy and isotope stratigraphy

9.4.3 Numerical time units (geochronology)

Chemostratiography and isotope stratigraphy are based on changing patterns of chemicals and isotopes through time (Chapter 7). There seems no reason why these and other methods do not have their own formal stratigraphy (chemozones and isostages) as do biostratigraphy and magnetostratigraphy. Sometimes, such as in the Upper Permian, chemical and isotopic stratigraphy is more precise than either biostratigraphy or magnetostratigraphy.

Geochronology gives ages in years. Geochronologic units are numerical divisions of time. They are based on units with the same time value, such as years, which can be counted from a known datum (usually the present). The only practical methods are based on either simple counting of annual varves or tree rings backwards, or on radiometric dating of various types (Chapter 7). Especially in the Phanerozoic, many geochronologic units simply date the time span of established

204 Chapter 9 Table 9.6 Currently recognized Phanerozoic stratigraphic system.

Mesozoic

Millions of years

Author, type locality, and reference-areas in North America

Quaternary

Recent Pleistocene

Last 10,000 Years 1.65 1.65

Jules Desnoyers proposed the Quaternary in 1829 for young deposits in France.

Tertiary

Pliocene Miocene Oligocene Eocene Paleocene

3.5 19 10 21 10

Giovanni Arduino deﬁned Tertiary rocks in 1760 in Italy as separate from the Primary and Secondary. Epochs were proposed by Charles Lyell in 1883 from type sections in France. Tertiary rocks are extensive in the High Plains and Coastal Plains in North America.

Neogene

Epoch

Paleogene

Cenozoic

Era Period

Cretaceous

80

Jurassic

55

Triassic

57

Permian

45

5 24 34 55 65

145

200

Pennsylvanian

Paleozoic

Mississippian

Carboniferous

251

300 11 311 44 355

Devonian

63

Silurian

22

Ordovician

50

Cambrian

55

418

440

490

545 Precambrian

3500

4,000

d’Omalius d’Halloy proposed the term Cretaceous in 1822 for the chalk-bearing sandstones and shales around the Paris Basin in France. Cretaceous rocks are well developed in the Coastal Plains and western interior of Canada and the United States. Alexander von Humbolt named the Jurassic from rocks in the Jura Mountains in northern Switzerland in 1795. Leopold von Buch proposed it as a system in 1839. Much scenery of the Colorado Plateau is carved into Jurassic rocks. Friedrich von Alberti introduced the term Triassic in 1834 for the threefold subdivision of salt-bearing rocks in northern Germany. Triassic rocks are extensive in the Mountain West. Roderick Impey Murchison named the Permian System from the province of Perm on the west ﬂank of the Ural Mountains in Russia. Reef-associated exposures in West Texas and New Mexico are standards for North America. Henry Shaler Williams proposed the Pennsylvanian System in 1891 for the coal-bearing rocks of Pennsylvania, equivalent to the Upper Carboniferous of Britain. Alexander Winchell proposed the Mississippian System in 1870 for the Lower Carboniferous rocks exposed in the Mississippi River Valley. Roderick Impey Murchison and Adam Sedgwick proposed the Devonian System for rocks in Devonshire, England. Devonian rocks are classically developed in New York and Pennsylvania. Roderick Impey Murchison proposed the Silurian System in 1835 for exposures in southern Wales. Silurian rocks are well developed in the Niagara Gorge region of New York and Ontario. Charles Lapworth named the Ordovician system in 1879 for a series of rocks involved in a controversy concerning the boundary of the the Cambrian and Silurian Systems. Ordovician rocks are well developed in New York, Ohio, Utah, and Nevada. Adam Sedgwick proposed the Cambrian System in 1835 for a series of shale and sandstone in northern Wales. Excellent exposures of Cambrian rocks occur in California, Utah, Alberta, British Columbia, and Wisconsin. Precambrian rocks are exposed in the cores of continents. In North America the most extensive exposures of Precambrian rocks are in the Canadian Shield surrounding Hudson Bay. They are also exposed in cores of the Rocky Mountains and Appalachian Mountains. Grand Canyon of the Colorado River exposes classic sequences. Precambrian rocks are essentially unfossiliferous.

Stratigraphic systems 205 chronostratigraphic units – which is simply the calibration of chronostratigraphic units. Such geochronologic units are given different names to the relative time units, though they roughly correspond (Table 9.5). This is useful since the relative time units are essentially ﬁxed (though imprecise), but the time units are constantly changing. For example, the Cretaceous System is ﬁxed by biostratigraphy, as is the Hauterivian Stage; but the Cretaceous Period and Hauterivian Age expand and contract and change positions like a jellyﬁsh (Fig. 9.11).

In the Precambrian, the divisions are primarily geochronologic in the absence of any correlatable relative methods of dating. The currently recognized Phanerozoic stratigraphic system and its calibration are shown in Table 9.6 and Appendix 3. “Basically there are only two concepts – rocks and time – with the rest just an obfuscation of the nomenclature.” (Ager (1973))

III Interpreting geologic history

Mt. Muztag Ata (7,500 meters) and tombs, Western Kun Lun, China.

Finding out how stratigraphic successions evolved and what controlled this evolution have always been the most interesting parts of stratigraphy. The four fundamental controls on the evolution of stratigraphic sequences on earth are tectonics, sea-level changes, climate, and biology. These are partly controlled by processes within the earth, such as plate tectonics and heat ﬂow; and partly by processes outside the earth, such as meteorite impacts and solar heat ﬂow

variations. Most changes in stratigraphic successions on earth involve complex combinations of many factors. Controls on other planets (such as Mars, which presently has no plate tectonics, oceans, or biology) may be more like the controls on early earth development, with magmatism and impacts still dominant (Patrick & Howe 1994), and may help us to interpret Archean stratigraphy.

10 Tectonics

10.1 10.2 10.3 10.4 10.5 10.6 10.7

Geodesy Hypsometry Gravity Isostasy Tectonics and sedimentary basins Exotic terranes Terrane analysis of orogenic belts

Tectonics studies the process, structures, and landforms associated with deformation. Understanding the tectonic controls on any planet requires knowing its GEODESY (shape and size), its HYPSOMETRY (topography), and how GRAVITY determines both. Gravitational and rotational effects on different densities of rocks and liquids determine ISOSTASY (which means “equal standing” in Greek) and explain the altitude and equilibrium of different parts of a planet. Vertical movements may adequately explain some CRATONIC BASINS, but most basins involve the horizontal movements of plate tectonics as well. Horizontal movements are orders of magnitude greater and faster than vertical movements, and form the majority of basin types recognized. Plate tectonics not only affects vertical changes, but also forms, transports, and juxtaposes EXOTIC TERRANES (stratigraphic successions that evolved independently, far away). The effects of all these

factors can be shown in analyses of mountain belts.

10.1 Geodesy Geodesy, the study of the shape and size of the earth, has a long history, extending back to the ancient Greeks of the 5th century bc, of whom Eratosthenes is best known. The link between the shape of the earth and gravity was ﬁrst demonstrated by Isaac Newton in the 17th century, who noted that a pendulum clock ran more slowly at the equator than in Paris, France. The surface at the equator was thus farther from the center of the earth than Paris was. The outward centrifugal force of the earth’s rotation must be balanced by increasing mass from poles to equator. The ideal shape of a uniform earth is a sphere ﬂattened by about 1/300, or

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210 Chapter 10

Fig. 10.1 How the hypsometric curve is calculated.

an oblate spheroid, with its equatorial diameter being about 43 km greater than its polar diameter. This REFERENCE SPHEROID (or ellipsoid) allows for increasing density with depth, but not for lateral variations in density and topography, which are obvious on the earth and the subject of hypsometry.

10.2 Hypsometry Hypsometry is the measurement and analysis of the distribution of topography. The HYPSOMETRIC CURVE describes the distribution of elevations across an area, from the scale of a planet to individual drainage basins. The curve is obtained by ﬁnding the areas between different heights (contour lines) and plotting the relative heights against the cumulative relative areas (Fig. 10.1). Drainage basins of different sizes can be compared since the curve is independent of size. A HYPSOMETRIC INTEGRAL (mean elevation - minimum elevation/maximum elevation - minimum elevation) describes the shape of the curve, and gives an estimate of “youthful” (deep incision and rugged relief), “mature” (intermediate), and “very mature” (shallow incision and subdued relief). Such concepts can be used to compare the development of sedimentary basins, if successive time slices of strata are used instead of relief (see Hay et al. 1988). Present earth hypsometry depends on segregation of continental and oceanic lithosphere, the rate of plate motions and the number of plates, and the base level of erosion to sea level (Fig. 10.2). This hypsometry may not apply to the past, for example in the late Paleozoic, when there were few plates and large supercontinents. The unimodal hypsometry of Venus indicates

that there is no similar continental–oceanic lithosphere division there (Keller & Pinter 1996).

10.3 Gravity Both geodesy and hypsometry are related to gravity. Gravity is measured in milligals, or thousandths of a gal (a GAL, after Galilieo, is equal to an acceleration of 1 cm/s2) because gravity differences over the earth are so minute. A GEOPOTENTIAL at a point on a planet’s surface is the combined value of the gravitational attraction and the rotational acceleration. For gravitational equilibrium, a planet’s surface should show a constant geopotential (after corrections for heights, etc.) based on its mass and rotation. This geopotential surface, or GEOID, differs from the reference spheroid, is perpendicular to plumb lines, and is the average sea level over the oceans and in imaginary canals below continents (Fig. 10.3). The geoid deviates from the reference spheroid by about 100 m or less; but the deviations (unlike the assumptions of Fig. 10.3) are independent of continents and ocean basins, and must be caused by density variations at great depths within the earth (Fig. 10.4). Reference spheroid, geoid, and topography never correspond, and elevation (orthometric height) is really height above the geoid. From a spherical perspective, the Amazon and Mississippi rivers ﬂow uphill. Gravity measurements need to be corrected to the reference spheroid for comparison (Telford et al. 1990). The ﬁrst correction is simply for latitude. The second, free air correction, corrects for height variations between the centers of mass of the instrument and the earth. This does not include the mass of any material

Tectonics 211

Fig. 10.2 Global hypsometric curve: the present land surface accounts for 29% of the earth’s surface area, but the break in slope on the hypsometric curve, the edge of the continental shelf, gives 37% as the true value for continental crust area (from Allen 1997, ﬁg. 1.37). (Reproduced with permission of Blackwell Publishing Ltd.)

Fig. 10.3 Exaggerated cross-section showing relationships between the reference spheroid (ellipsoid) and the geoid under continents and ocean basins (after Heiskanen & Vening Meinesz 1958, ﬁg. 8.5).

between the station and the geoid, so a third correction, the Bouger correction, is required to correct for the effects of mass lying above the spheroid (Fig. 10.5a). The correction assumes an average-thickness inﬁnite slab – not true, but accurate enough for the purpose (Fig. 10.5b).

The Bouger correction also assumes that the rock material has a uniform density throughout, which of course is not always the case. Furthermore, in areas of signiﬁcant differential relief, additional corrections need to be made for local topography. The difference between observed and predicted anomalies (based on the

212 Chapter 10

Fig. 10.4 The geoid. Contours are in meters, at 10 m intervals (from Smith 1973, ﬁg. 2.5.1). (Reproduced with permission of The MIT Press.)

Fig. 10.5 (a) The Bouger correction for the mass between the station and the geoid, assuming (b) an inﬁnite slab of thickness h and density p (from Lillie 1999, ﬁg. 8.8).

reference spheroid) is known as a GRAVITY ANOMALY. BOUGER GRAVITY ANOMALIES are strongly negative in mountain areas because the correction does not include their low-density roots required for isostatic compensation

(see Section 10.4). Gravity anomalies are extremely useful in determining the large-scale mass distributions within a sedimentary basin, but there are usually several basin-ﬁll models that ﬁt the gravity data. So, seismic

Tectonics 213

Fig. 10.6 Three possible structural/density models that satisfy the Bouger anomaly calculations across the Mid-Atlantic Ridge. Case I has uniform intermediate (anomalous) mantle (p = 3.15). Case II has intermediate mantle increasing in density downwards. Case III assumes low density near the ridge axis. Density units are 103 kg/m3 (from Smith 1973, ﬁg. 2.9.14). (Reproduced with permission of The MIT Press.)

analysis may help in choosing the most likely model: for example, in Fig. 10.6, seismic analysis makes case III the most likely structural/density model for the MidAtlantic Ridge. Even here, the low-density plume now inferred for the mid-oceanic ridges is not included in this pre-plate tectonic example. On a small scale, gravity varies because of near-surface changes in rock density. They are negative when there is a mass deﬁcit, and positive when there is a mass excess. Such local gravity anomalies can be used to

detect economic mineral deposits. A dense ore body gives positive gravity anomalies, whereas a thicker sedimentary column gives negative gravity anomalies.

10.4 Isostasy Isostasy is a state of “ﬂoating” equilibrium; that is, a planet’s surface should have the same gravitational potential everywhere. Common analogies are blocks of

214 Chapter 10

Fig. 10.7 Different icebergs ﬂoat with the same relative proportions above and below water.

wood or ice ﬂoating in water (Fig. 10.7). Water is denser than wood and ice. So wood, ice, and water have different volumes for the same mass. d ice ¥ Vice = d wood ¥ Vwood = d water ¥ Vwater

(10.1)

For freshwater at 5 °C, ice at 0 °C, and your average softwood, typical density values are 1.0 (freshwater), 0.9 (ice), and 0.5 (softwood) g/cm3. So, to get the same mass as 1.0 m3 of water, you need 1.11 m3 of ice, and 0.11 m3 will be above the water. For softwood, you need 2.0 m3, with 1.0 m3 above the water. This is the basis of Archimedes’ principle that a ﬂoating body displaces its own weight of water. The ﬂoating body exerts a force (mass ¥ acceleration of gravity) equal to that of the displaced water. Water ﬂows under, or away from, the ﬂoating ice or wood until they weigh the same as the displaced water. This is a more useful way of looking at the problem than considering the height or thickness of the blocks, as is frequently done for simplicity, because few ﬂoating objects are regular cubes. (The air displaced by the solid above water also subtracts, marginally, from its buoyancy.) The volume of ﬂoating blocks (here ice and wood) above an ambient ﬂuid (here water) is proportional to the relative densities of ﬂoating block and ﬂuid: volume solid (above fluid) = total volume solid (density fluid -density solid )va s = vt s (d w - d s )

(10.2)

or va s vt s = (d w - d s )

(10.3)

From the densities given above, ice at 0 °C will ﬂoat with one-tenth of its volume above water, while softwood will ﬂoat with half its volume above water. Rafts of softwood ﬂoat higher than rafts of hardwood (but they become denser by soaking up water faster, and then sink – as the balsa raft did on the Kon Tiki Expedition). This calculation assumes that the material has a uniform density (AIREY ISOSTASY), but if materials vary in density, then they can ﬂoat at different levels without roots (PRATT ISOSTASY). Since we now know that topography is supported by subcrustal roots, Airey isostasy is generally used, but simple physics indicates that rocks must increase in density with depth due to conﬁning pressure, so Pratt isostasy must apply to some extent. Normally, different sections of the lithosphere are given different average densities to overcome this, often based on seismic velocities (Fig. 10.8). Isostasy can be changed by changing the relative densities of the ﬂoating object and the supporting medium (or both together), or by changing the volume of the ﬂoating object. An iceberg with stones will be denser and ﬂoat lower than one without stones, and it will ﬂoat slightly higher in sea water (density 1.03 g/cm3) than in freshwater (density 1.00 g/cm3). These density changes are due to changes in composition. A density change without a composition change occurs when materials are heated or cooled: an iceberg in colder (hence denser) water will ﬂoat slightly higher than one in warmer water. Changes in volume affect the height of the “ﬂoating” object: a melting iceberg will decrease in average height (though its relative proportions above and below water stay the same). The LITHOSPHERE (the rigid outer part of the earth),

Tectonics 215

Fig. 10.8 Various seismic velocity sections in the western USA (from Smith 1973, ﬁg. 2.9.8). (Reproduced with permission of The MIT Press.)

Fig. 10.9 Loading of the lithosphere with an ice sheet causes asthenosphere ﬂow and forebulge development due to ﬂexural rigidity of the lithosphere.

analogous to the wood or iceberg, is in ﬂoating equilibrium with the ASTHENOSPHERE (the plastic upper mantle of the earth), analogous to the water. However, the lithosphere has some elasticity or ﬂexural rigidity, while the asthenosphere behaves basically as a plastic. If the lithosphere is loaded or unloaded with extra weight, such as a mountain belt, sediment, an ice sheet, a lake, or more water in the sea, it ﬁrst responds elastically, and then takes some time to respond plastically by ﬂow of asthenosphere (Fig. 10.9). The time taken (and the plasticity, or rheology, of the mantle) can be estimated from the isostatic rebound of Canada and Scandinavia after melting of the last Pleistocene ice-caps. Mantle ﬂow re-established isostatic equilibrium almost entirely within 1000 to 10,000

years; that is, the response time for isostatic equilibrium is well within the precision of dating in pre-Quaternary times, and is essentially instantaneous over geologic time. Furthermore, the isostatic depression of polar regions, and consequent ﬂow of asthenosphere toward the equator (and vice versa) during glacial (interglacial) periods, has implications not only for relative sea level, but also for the rotational stability of a planet (Fig. 10.10) (Peltier 1998). Loads can be supported non-isostatically over variable periods of time, depending on the ﬂexural rigidity of the lithosphere, which varies with the nature of the lithosphere, the wavelength of the load, and the amount and duration of the applied stress (Fig. 10.11). This is like standing on a diving board: the board ﬂexes but supports your weight depending on how thick it is, what it is made of, how heavy you are, and whether you are lying ﬂat or upright. For loads of sufﬁciently long wavelength, such as large mountain belts, the lithosphere has no rigidity and the load is fully compensated (in hydrostatic equilibrium). The degree of compensation depends both on the wavelength and the duration of the load (Allen 1997). Flexural loading by ice-caps, volcanoes, and overthrusting leads to the formation of marginal depressions and peripheral FOREBULGES. These gradually migrate through time as the lithosphere deforms plastically, depending on the rate and extent of convergence (Fig. 10.12). It is

216 Chapter 10

Fig. 10.10 Radiocarbon dates for the period 5000–2500 years BP related to sea-level. Note the zone of submergence around regions of known rapid postglacial uplift (from Walcott 1972). (Copyright (1972), with permission from Elsevier Science.)

Fig. 10.11 Flexural rigidity. Thick (continental) lithosphere supports a volcano better (e.g. Mount St Helens) and has wider, shallower marginal basins than thin (oceanic) lithosphere, which has narrower, deeper marginal basins (e.g. Hawaii). Very thin lithosphere collapses into local isostatic equilibrium under volcano load (no present-day earth examples, but perhaps some Archean volcanic piles behaved like this).

worth emphasizing that the depression depths are only a small fraction of the height of the load. The effect of this on stratigraphic successions is to cause differential relative changes of base level, which migrate through time and mimic eustatic changes of sea level.

Simple isostatic movements involve changes in gravitational potential due to thickening and thinning, or heating and cooling of continental or oceanic crust; sedimentary loading and erosional unloading of the earth’s surface (Davies 1999).

Tectonics 217 10.4.1 Uplift

10.4.2 Subsidence

Large-scale uplift requires an area to become lighter, either by a reduction in density or an increase in thickness of a lighter material. A reduction in density may be caused by heating or by insertion of less dense material. Mid-oceanic rises and continental rifts lie on swells caused by a reduction in density of the upper mantle by heating. Thickening of less dense continental material, by squeezing in magmatic arcs and underthrusting in continental collision mountains, together with heating and intrusion of less dense “granitic” intrusions, causes mountains to rise (Fig. 10.13).

Large-scale subsidence requires an area to become heavier, either by an increase in density or by an increase in the thickness of denser material. An increase in density may be caused by cooling or by insertion of denser material. Oceanic lithosphere cools exponentially and becomes thicker and denser away from a mid-oceanic rise; in fact, the age of an ocean ﬂoor can be inferred from its depth (Fig. 10.14). Sagging of the central rifts on continental swells is partly caused by the eruption and intrusion of heavier basaltic magmas, lavas, and dykes in the continental crust. 10.4.3 Lateral ﬂow Supporting isostatic roots developed tectonically are partly removed by lateral ﬂow at depth. Many intermontane basins are preserved because these have been let down to sea level by subcrustal ﬂow (Fig. 10.15). High mountains require continuing compression to sustain them, and subside more by lateral ﬂow than by surface erosion (Dewey 1988). 10.4.4 Local uplift and subsidence

Fig. 10.12 Migrating forebulge and a narrowing and deepening basin develop after loading by volcanoes, thrust sheets, deltas, or other restricted loads (from Leeder 1999, ﬁg. 27.9). (Reproduced with permission of Blackwell Publishing Ltd.)

Fig. 10.13 Uplift by thickening of continental crust, with heating and intrusion, Andes (23°S) (from Fowler 1990, ﬁg. 8.47). (Reproduced with permission of Cambridge University Press.)

Relatively small areas are frequently not in isostatic equilibrium. Thrust loading forms more depressed foredeeps and higher adjacent mountains (Fig. 10.12). The

218 Chapter 10

Fig. 10.14 Thickening and cooling of oceanic lithosphere (from Howell 1989, ﬁg. 2.7). (With kind permission of Kluwer Academic Publishers.)

(a)

Fig. 10.15 Lateral ﬂow and subsidence of crustal material (from Kearey & Vine 1996, ﬁg. 10.31). (Reproduced with permission of Blackwell Publishing Ltd.)

(b) Amount removed by erosion

Himalaya

Tibet

Continental crust

Fig. 10.16 (a) Central Himalaya from the east (courtesy NASA). (b) Cross-section showing how valley erosion causes isostatic uplift of the peaks between them.

extreme height of some Himalayan peaks is caused by isostatic uplift as valleys are cut between them (Fig. 10.16). The highest mountains are always surrounded by the deepest valleys (Fig. 10.17).

10.5 Tectonics and sedimentary basins Tectonics fundamentally controls the formation and development of sedimentary basins, which tend to be of

two basic types: those formed by vertical tectonics apparently unrelated to plate motions, and those formed by plate movements and associated deformation (Table 10.1). 10.5.1 Basins formed by vertical movements There are two main types of basin formed by vertical movements: impact craters and basins, and cratonic basins.

Tectonics 219

Fig. 10.17 Northern Apennines, Italy: gravitational collapse and differential erosion (courtesy F. Ghisetti, Catania).

Table 10.1 Tectonic basin classiﬁcation (modiﬁed from Einsele 1992, table 1.1). Basin category

Special basin type or synomym(s)

Underlying crust

Style of tectonics

Basin characteristics

Cratonic basins

Epicontinental basins, infracratonic basins Graben structures, rift valleys and rift zones, aulacogens

Continental

Divergence

Large areas, slow subsidence

Continental

Divergence

Transitional

Divergence + shear

Oceanic

Divergence

Oceanic

Convergence

Relatively narrow basins, fault-bounded rapid subsidence during early rifting Asymmetric basins partly outbuilding of sediment, moderate to low subsidence during later stages Large, asymmetric, slow subsidence Partly asymmetric, greatly varying depth and subsidence

Continental rift basins

Passive margin basins

Oceanic basins Subduction-related basins Magmatic arc basins

Continental collision basins

Strike-slip basins

Tensional-rifted basins, tension-sheared basins, sunk margin basins Nascent ocean basin (growing oceanic basin) Deep-sea trenches Fore-arc basins, backarc basins, inter-arc basins Remnant basins

Transitional, oceanic Dominantly divergence Oceanic

Convergence

Foreland basins (peripheral), Continental retro-arc basins (intramontane), broken foreland basins Terrane-related basins Oceanic

Crustal ﬂexuring, local convergence or transform motions

Pull-apart basins (transtensional) and transpressional basins

Transform motion, ± divergence or convergence

Continental and/or oceanic

Activated subsidence due to rapid sedimentary loading Asymmetric basins, trend to increasing subsidence, uplift, and subsidence Similar to back-arc basins Relatively small, elongate, rapid subsidence

220 Chapter 10

Fig. 10.18 Nordlinger church built of impact breccia.

rings are melted, brecciated, and shocked materials (Fig. 3.4b). Impact basins are wider than 180–300 km on the Moon, but smaller on larger bodies, and have multiple rings and central uplifts. They mostly formed by huge impacts early in planetary history. Impact basins have been found on the earth, though the lower size limit is not yet known, and most can only be remotely imaged below covering sediments. The Chicxulub lateCretaceous impact basin, covered by younger limestone, has an outer ring 300 km in diameter, with inner rings at 150 and 104 km (Fig. 10.19). Some of the large, circular cratonic basins, such as the Hudson Bay and Michigan basins, discussed below, may have started as large impact craters. In fact, it is difﬁcult to envisage any other cause for the perfect half-circle of eastern Hudson Bay (which becomes two-thirds of a circle with younger rifting removed; Fig. 10.20), although no good evidence for impact has been found so far. Post-impact sedimentation in a crater depends on other factors, such as plate tectonics and climate. 10.5.1.2 Cratonic basins

10.5.1.1 Impact craters and basins IMPACT CRATERS and their deposits are usually ignored in

stratigraphy and basin analysis texts. However, the oldest lunar (and other inner solar system) surfaces are saturated with impact craters produced by an intense bombardment of the inner solar system from about 4.6 to 3.9 billion years ago. This must have affected the earth and has implications for early Archean basin formation and development, which have not yet been sufﬁciently investigated (Melosh 1990; Ringwood 1990). Impact craters are classed as simple and complex (see Section 3.2.1). Simple craters form heavily fractured and brecciated bowls, surrounded by ejecta and partially ﬁlled with breccia and impact meltrock lenses (Fig. 3.4a). Earth has many recognizable simple craters, though most are degraded or obscured by overlying sediment. One of the most interesting small complex craters (at least for beer drinkers) is the 20 km diameter Ries Crater in southern Germany (Fig. 10.18). Complex craters on earth start at crater diameters of between 2 and 4 km, and on the Moon, with its lower gravity, between 15 and 20 km. Complex craters can be one hundred times wider than they are deep. A central uplift is surrounded by rings marking the collapse of transient earth waves, while between the uplift and the

CRATONIC BASINS occur within continents, tend to be bowlshaped, and have long histories of generally slow but sporadic subsidence. They are ﬁlled with thin but extensive sedimentary units deposited near sea level, which often show a very regular thickening towards the basin center (a “bulls-eye” pattern) (Fig. 10.21). While the total sediment within them is controlled by their slow subsidence, each individual unconformity-bounded unit is controlled by sea-level changes. The origin of cratonic basins has been a puzzle for a long time. Their slow subsidence obviously requires some gradual increase in density beneath them; but how this occurs is debatable – cooling of mantle plumes or phase changes at the base of the lithosphere? The circularity of some suggests crustal or lithospheric ring fracturing due to either internal or external shocks. Some may originate as extraterrestrial impacts, as at least some have central rifted uplifts (Leighton & Kolata 1990).

10.5.2 Plate tectonics and basin development Plate tectonics fundamentally determines the size, shape, and development of all basins apart from impact and cratonic basins. The largest basins with the thickest sediments are at the junctions of continents and oceans, where rivers deliver huge volumes of sediment to

Tectonics 221 (a)

Fig. 10.19 Chicxulub impact basin, Yyucatan, Mexico: (a) gravity anomalies showing multiple rings (courtesy Geological Survey of Canada); (b) seismic cross-section with post-impact basins (courtesy P. Claes, Berlin).

deep-water areas capable of accommodating and preserving them, and these are the trailing-edge subsiding passive margins of continents formed by ocean opening (Fig. 10.22). Recent plate-tectonic basin form and development, and synthetic models based on them, can be used to interpret and classify preserved ancient sedimentary basins (Allen & Allen 1990; Busby & Ingersoll 1995; Miall 2000). A characteristic scheme of basin evolution can be related to the standard opening and closing ocean models of plate tectonics (Fig. 10.23). This evolu-

tion can be short-circuited at any stage, and lateral movements along transform faults can juxtapose basins that evolved in unrelated tectonic areas. Furthermore, the actual environmental complexes developed are also dependent on climate; so that a rift in a polar tundra will have a very different sediment ﬁll to a rift in a tropical desert. Nevertheless, the simple basin models described below allow various environmental complexes to be related in a genetic way to show the evolution of an area. The various basin types in the various plate tectonic environments (and not quite the same as Table 10.1) are

222 Chapter 10 (a)

Fig. 10.20 (a) Shuttle view of the eastern arc of Hudson Bay, radius 230 km (courtesy NASA); (b) two-thirds of a circle restored by closing James Bay (from Goodings & Brookﬁeld 1992, ﬁg. 10). (Copyright (1992), with permission from Elsevier Science.)

divided into those of continental rifts; mid-oceanic rifts; ocean ﬂoors; passive margins; subduction zones; island arcs; Andean belts; collision orogens; and transform belts. Each of these is brieﬂy summarized in turn. Only recent or Cenozoic examples are included as older ones have to be interpreted. 10.5.2.1 Continental rifts CONTINENTAL RIFTS form anywhere where continental crust is being pulled apart, which can occur in many tectonic situations (equating continental rifts with the start of continental splitting and ocean formation is unjustiﬁed). Continental rifts occur above rising mantle plumes and rolls beneath continents (e.g. East Africa), at fault junctions along transcurrent faults (e.g. eastern Asia), and as collapse structures above gravitationally spreading crust (e.g. Tibet, Andes). The east African rifts show typical features (Fig. 10.24). Physical and time

correlation is difﬁcult because of rapid lithofacies and (endemic) biofacies changes; while endemism makes correlation with standard biostratigraphic units very difﬁcult. Large rift lakes isolated for considerable time periods may have very peculiar, entirely endemic, faunas; for example, the Devonian lakes of northern Scotland. Rifts that mark the start of splitting continents evolve ﬁrst into rifts ﬂooded by the sea and then into open oceans with mid-oceanic rifts. The transitional stage is marked by isolated ﬂooded rifts, such as the Red Sea and the Gulf of Aden (Fig. 10.24a). Ancient failed rift arms are called aulacogens (Fig. 10.24b,c). Rifts typically have the following characteristics (Fig. 10.25): 䊏 narrow, fault-bounded sedimentary basins arranged en echelon; 䊏 successions disrupted by magmatism and faulting; 䊏 basaltic and peculiar alkaline magmatism due to ponding and differentiation of magma;

Tectonics 223

Fig. 10.21 Cratonic basins of the Hudson Bay area (from Norris 1986, ﬁg. 2.5).

䊏 contact metamorphism around intrusions; 䊏 normal and minor strike-slip faulting; 䊏 rapid and abrupt facies changes, both horizontally and vertically; 䊏 sediment ﬁlls depending on climate – evaporitic ﬁlls in dry climates, lacustrine ﬁlls in wet climates. 10.5.2.2 Mid-oceanic rifts MID-OCEANIC RIFTS resemble continental rifts, except that they are underwater; chemical reactions of hot circulating brines cause hydrothermal metamorphism of the basalt lavas; coarse clastic sediments are less common; and biochemical sediments (such as foraminiferal oozes and radiolarian cherts) are more abundant (Fig. 10.26).

10.5.2.3 Ocean ﬂoors OCEAN FLOORS consist of ophiolitic basement overlain by a variety of volcanic and sediment types. Hot-spot basaltic shield volcanoes (e.g. Hawaii) and plateaus form seamounts, which may choke subduction zones. Thin chemical sediments, such as carbonate and silica oozes, form in open-ocean environments and change into limestone and chert respectively. Shales and graded volcanic ashes may be dominant near magmatic arcs. Mature graded sandstones increase in thickness and proportion towards passive continental margins, where large chaotic slides may occur. Physical and time correlation is often relatively easy because of widespread and interﬁngering lithofacies, distinct event

224 Chapter 10

Fig. 10.22 Sediment thickness in modern basins (from Blatt 1982, ﬁg. 1.6).

Fig. 10.23 Wilson cycle with successive basin formation (from Kearey & Vine 1996, ﬁg. 10.11). (Reproduced with permission of Blackwell Publishing Ltd.)

Tectonics 225 (b)

Fig. 10.24 (a) Variety of rifts and narrow ocean basins, northeast Africa (from Kearey & Vine 1996, ﬁg. 10.21). (Reproduced with permission of Blackwell Publishing Ltd.) (b) Environmental mosaic in the Afar triangle (courtesy NASA). (c) Failed rifts (aulacogens) in the southern USA (from Einsele 1992, ﬁg. 1.1). (Copyright (1992). Reproduced with permission of SpringerVerlag.)

layers (such as identiﬁable and datable ashes), and widespread zone fossils of pelagic and nektonic organisms. Nevertheless, different parts of the same ocean may have distinct faunal provinces due to oceanographic conditions. Ocean ﬂoors typically have the following characteristics (Fig. 10.27): 䊏 extensive basins of low relief, which may be offset across transform faults;

䊏 thin successions dominated by chemical, biogenic, and ﬁne-grained clastic sediments; 䊏 simple basaltic volcanism, often hydrothermally altered; 䊏 local contact and hydrothermal metamorphism; 䊏 normal and minor strike-slip faulting, including the special case of transform faults; 䊏 subdued and gradual horizontal facies changes, although vertical changes may be sharp;

226 Chapter 10 (a)

Fig. 10.25 Rift valley ﬁlls: (a) wet alluvial-lake ﬁll at the head of the Gulf of California, with the Colorado delta leading into it, and the Sonoran Desert adjacent to it; (b) dry evaporitic ﬁll of rift (from Mitchell & Reading 1986, part of ﬁg. 149). (Reproduced with permission of Blackwell Publishing Ltd.)

Fig. 10.26 Characteristics of mid-oceanic rifts (from Kearey & Vine 1996, ﬁg. 12.56). (Reproduced with permission of Blackwell Publishing Ltd.)

䊏 sediment ﬁlls depending on ocean conditions and proximity to other environments; 䊏 easily correlatable successions based on planktonic and nektonic organisms. 10.5.2.4 Passive margins PASSIVE MARGINS consist of slowly subsiding, thinned continental crust at the edge of an opening ocean.

Sedimentation on the margins is mainly controlled by thermal subsidence, eustatic sea-level changes, and climate. Sediments consist of widespread, relatively uniform deposits (though showing local mosaics) which grade gradually into one another; however, submarine canyons may transport nearshore sediments across the margin to form deep-water fans (Fig. 10.28). In humid areas, mature clastic sediment is often supplied by rivers ﬂowing long

Fig. 10.27 Ocean-ﬂoor characteristics (from Press & Siever 1986, ﬁg. 11.39).

Fig. 10.28 Block diagram of a passive margin (from Moores & Twiss 1995, ﬁg. 3.17). (© (1995) by W.H. Freeman and Company. Used with permission.)

228 Chapter 10

Fig. 10.29 Passive-margin cross-section, eastern USA (from Mitchell & Reading 1986, ﬁg. 14.21). (Reproduced with permission of Blackwell Publishing Ltd.)

distances from magmatic arcs on the other side of the continent (e.g. North America, South America). Dry areas may have passive margins dominated by carbonates. Passive margins typically have the following characteristics (Fig. 10.29): 䊏 extensive basins of low relief, which usually deepen oceanwards; 䊏 variable thin (nearshore) to thick (offshore) successions; 䊏 no contemporary volcanic rocks; 䊏 no metamorphism; 䊏 little or no contemporary deformation; 䊏 subdued and gradual facies changes, often with a consistent nearshore–offshore pattern; 䊏 sediment ﬁlls depending on relative sea level, climate, and sediment supply; 䊏 successions that are often correlatable on zone fossils, but also show strong biofacies control. Passive margin successions are common in the stratigraphic record because they accumulate a very large proportion of deposited sediment. Many facies models based on modern passive margins have been erected to

interpret these ancient deposits (see Walker & James 1992; Reading 1996). Passive margins crossing many degrees of latitude show great changes due to the climatic controls on extra- and intra-basinal sedimentation (Fig. 10.30). 10.5.2.5 Subduction zones SUBDUCTION ZONES develop where thickened oceanic lithosphere starts gravitationally sinking, plunging below other oceanic lithosphere or the edge of a continent to form deep ocean trenches. Sediments are marine, of variable thickness, and usually easily correlatable with zone fossils. Island arcs or Andean belt volcanoes develop around 100–400 km behind the trench, depending on the angle of incidence of the subducting slab. Subduction zones have the following characteristics (Fig. 10.31): 䊏 narrow and deep linear basins; 䊏 variably thin to thick marine sediment ﬁlls dominated by graded turbidity current deposits, depending on adjacent areas;

Tectonics 229

Fig. 10.30 Facies changes down the eastern North American passive margin (from Milliman et al. 1972, ﬁg. 13). (Copyright © 1972. Reproduced with permission of the publisher, the Geological Society of America, Boulder, Colorado, USA. ) Contours in meters.

Fig. 10.31 Character of subduction zone basins (trenches) (from Pickering et al. 1989, ﬁg. 11.12). (Reproduced with permission of HarperCollins.)

䊏 no contemporary volcanic rocks; 䊏 low-temperature, high-pressure metamorphism along faults; 䊏 contemporary deformation, with chaotic melanges related to faulting and slumping;

䊏 rapid facies changes across, but subdued facies changes along, the basins, since turbidity currents tend to ﬂow axially; 䊏 sediment ﬁlls related to the character of nearby environments, frequently volcaniclastic;

230 Chapter 10

Fig. 10.32 (a) Relationships of some island-arc basins (from Einsele 1992, ﬁg. 1.2) (Copyright (1992). Reproduced with permission of Springer-Verlag.); (b) back-arc basin sedimentation (from Bruhn & Dalziel 1977, ﬁg. 2a).

䊏 successions correlatable on zone fossils with few benthic fossils. 10.5.2.6 Island arcs ISLAND ARCS develop above oceanic crust behind subduction zones. This often happens at the edges of oceans, where the lithosphere tends to be oldest and thickest. Some arcs apparently develop by sucking of a continental marginal arc out into the ocean with the develop-

ment of a marginal basin behind (e.g. Japan). Island arc magmatism only starts when the subducting lithosphere reaches a depth of 200 km. At a subduction angle of 30°, and a subduction rate of 0.1 m/year, it thus takes around two million years for an island arc to begin developing after the start of subduction. Basin types vary from fore-arc marine to intra-arc continental. Island arc basins have the following characteristics (Fig. 10.32):

Tectonics 231

Fig. 10.33 Marginal (back-arc) basins of the western Paciﬁc Ocean. The Songliao Basin is a retro-arc basin (see Fig. 10.34).

䊏 variably wide (fore-arc) to narrow faulted (intra-arc) linear basins; 䊏 generally thick, mostly marine volcaniclastic ﬁlls; 䊏 contemporary basaltic to andesitic volcanic rocks; 䊏 moderate pressure Abukuma-type regional metamorphism in places; 䊏 strong contemporary deformation reverse, with normal, reverse, and strike-slip faulting; 䊏 rapid facies changes, with individual basin ﬁlls often distinctive; 䊏 dominantly volcaniclastic sediments, but some carbonates in shallow marine areas; 䊏 marine successions that are relatively easy to correlate with fossils.

Back-arc basins resemble ocean ﬂoors but with volcaniclastic input from the arc. Back-arc, or marginal, basins occupy vast areas of the western Paciﬁc Ocean (Fig. 10.33). 10.5.2.7 Andean belts ANDEAN BELTS are magmatic arcs that develop on continental crust. They are presently found only along the western side of Central and Southern America where the Atlantic plate is rapidly overriding younger portions of the Paciﬁc plate, and in Indonesia where the Southeast Asian plate is overriding the Australian plate. Some Andean belts have

232 Chapter 10

Fig. 10.34 (a) Relationships of some Andean basins (from Einsele 1992, ﬁg. 1.2) (Copyright (1992). Reproduced with permission of Springer-Verlag.); (b) Andean cross-section showing basin sedimentation (from Megard 1989, ﬁg. 14.12A). Used by permission of Oxford University Press, Inc.)

apparently converted into island arcs when the belts split off from a continent, with the development of marginal basins: good examples are the Japanese Islands, which separated from the mainland of China in the Oligocene. Andean belts have the following characteristics (Fig. 10.34): 䊏 variably wide (fore-arc) to narrow faulted (intra-arc) to wide overthrust (retro-arc) basins; 䊏 generally thick clastic ﬁlls, mostly non-marine in intra-arc, mixed in retro-arc basins; 䊏 contemporary andesitic to rhyolitic volcanic rocks, with granodiorite–granite intrusions; 䊏 high-pressure Barrovian regional metamorphism; 䊏 strong contemporary deformation such as island arcs;

䊏 rapid facies changes within individual basins; 䊏 mixed igneous, metamorphic, and basementderived sediments; 䊏 mixed marine–continental successions which may be difﬁcult to correlate. 10.5.2.8 Collision orogens COLLISION OROGENS are areas of extreme deformation where magmatic arcs collide with continents (as in western North America) or where two continents collide (as in the Himalaya and Tibet). Irregular continental margins form foreland basins (where one continent overrides another) and remnant basins (where oceanic crust survives). Overriding continents have high-altitude rifts and strike-slip basins

Tectonics 233

Fig. 10.35 (a) Relationships of collision orogen basins (from Einsele 1992, ﬁg. 1.2c). (Copyright (1992). Reproduced with permission of Springer-Verlag.) (b) Foreland basin cross-section, in which X shows successive positions of onlap for various times, T (from Kearey & Vine 1996, ﬁg. 10.40). (Reproduced by permission of Blackwell Publishing Ltd.)

caused by gravitational collapse and oblique collision. Individual basins are difﬁcult to correlate because of local tectonics, many unconformities, and endemic faunas.

Collision orogens have the following characteristics (Fig. 10.35): 䊏 generally narrow, faulted basins, with signiﬁcant growth folds in foreland basins;

234 Chapter 10 (a)

(b)

Fig. 10.36 (a) Relationships of transform basins (from Einsele 1992, ﬁg. 1.3) (Copyright (1992). Reproduced with permission of Springer-Verlag.); (b) transform basins, Gulf of California (from Crowell 1974, ﬁg. 2). (Reproduced with permission of SEPM,)

䊏 thick, clastic, dominantly non-marine ﬁlls which may begin with marine clastics; 䊏 no contemporary volcanism, but local anatectic granitic plutonism along thrusts; 䊏 high-pressure Barrovian regional metamorphism rapidly exhumed in the overriding continent; 䊏 strong contemporary deformation, with deformation migrating in foreland basins; 䊏 rapid facies changes within individual basins; 䊏 mixed basement and cover-derived sediments, with minor old arc-derived clastics; 䊏 marine to predominantly continental successions that are difﬁcult to correlate.

10.5.2.9 Transform belts TRANSFORM BELTS are areas of dominantly strike-slip faulting between plates, and have complex, narrow basins bounded by oblique-slip, normal, and reverse faults. They develop on both oceanic and continental lithosphere and are important in stratigraphy since transform faults can juxtapose units that evolved long distances apart, often without any signiﬁcant deformation to advertise the fact. In contrast, transform belts can also separate very distinct basin successions that evolved very close together. Transform belts have the following characteristics (Fig. 10.36):

Tectonics 235 䊏 narrow, faulted, and often very deep, basins; 䊏 thick marine (oceanic) or non-marine (continental) clastic ﬁlls; 䊏 no contemporary volcanic or plutonism; 䊏 no metamorphism, apart from local dislocation metamorphism along faults; 䊏 strong contemporary deformation, mostly normal and reverse oblique faulting; 䊏 rapid facies changes within and among individual basins; 䊏 sediment compositions depending on the local bedrock; 䊏 marine (oceanic) to continental successions. Distinctive features of transform basins are contemporary oblique-slip, normal, and reverse faults and “sourceless” sediments because lateral movement can separate sediments from their original sources. In California, three Cretaceous units were once part of the same alluvial–submarine fan system, although several hundred kilometers now separate one unit and its source (Fig. 10.37a). Adjacent en echelon submarine transform-bounded basins, such as rifts, may have very different successions depending on sediment supply and local tectonic effects, without movement (Fig. 10.37b).

set in a foliated matrix of clastics or serpentine, and are typical of subduction or transform-fault settings. Dispersion occurs during rifting or strike-slip (transform) faulting (Fig. 10.39). Transform faulting, in particular, may juxtapose very different, as well as separate distinct terranes. Amalgamation occurs during subduction, continental collision, and transform faulting. It can be recognized by provenance linking, overlap sequences, stitching plutons, and welding metamorphism (Fig. 10.40), and by the breakdown of marine biological provinces. Provenance linking has identiﬁable sources on one terrane providing sediment to adjacent terranes. Overlap sequences, stitching plutons, and welding metamorphism cut across the tectonic joins of terranes, and dating these determines the time of amalgamation. Good examples occur throughout western North America (see Monger 1984). Provincial linking is when originally divergent fossil assemblages become similar or identical, indicating that barriers to migration – often wide oceans – have broken down. A good example is the progressive reduction in Ordovician faunal provincialism in the North Atlantic during the closure of the Iapetus Ocean, culminating in the Caledonian collision orogeny. Recognizing terranes is an essential preliminary to the analysis of an orogenic belt.

10.6 Exotic terranes An EXOTIC TERRANE is a piece of an area that has some peculiar characteristics that set it apart from adjacent pieces. Most exotic terranes form parts of orogenic belts, and have been introduced by large lateral movements during the development of the mountain belt (Fig. 10.38). A more general term is tectonostratigraphic terrane, a fault-bounded rock unit with a distinct stratigraphy that is a unique piece in an orogenic collage (Howell 1995). Such terranes are normally distinguished by their stratigraphy, but rarely a metamorphic or structural overprint may be useful. So, there are three categories: stratigraphic, metamorphic, and disrupted terranes. Stratigraphic terranes may be: (i) displaced fragments of continents; (ii) fragments of continental margins; (iii) fragments of volcanic arcs; or (iv) fragments of ocean basins. Metamorphic terranes show recrystallization and terrane-wide penetrative metamorphism, which obliterates the original stratigraphy. Disrupted terranes have blocks of diverse rocks of different ages

10.7 Terrane analysis of orogenic belts Mountain belts (or orogens) have various tectonic units assembled after frequently complex individual histories. Ideally, one should start with recent units and then follow their individual histories backwards in time, as various pieces amalgamate, disperse, and fragment. Terrane analysis is a particularly suitable method, despite the criticisms of Sengor (1984). It is a natural outgrowth of plate-tectonic analysis and an extension of earlier collage concepts of orogenic evolution (Helwig 1974). Any orogen is a collage of different elements (terranes) with different histories. An essential ﬁrst step is to deﬁne the different terranes at different times. We may then, hopefully, follow their complex histories through time. Major terrane boundaries usually occur at sutures, marking the sites of former oceans, so all data must be carefully plotted on suture maps for them to be useful in reconstructions. The analysis requires a multidisciplinary approach involving regional stratigraphy and structure, petrology, paleobiogeography, palaeoclimatology,

(a)

(b)

Fig. 10.37 (a) Lateral removal of source by transform faulting; thick arrows show dispersed segments of a single depositional system (from Blatt & Tracy 1996, ﬁg. 13.4). (b) Transform-bounded basins off South America. (from Francheteau & Le Pichon 1972, ﬁg. 11).

Tectonics 237

Fig. 10.38 Terranes of the western North American Cordillera (from Saleeby 1983, ﬁg. 1).

Fig. 10.39 Dispersion of three amalgamated terranes by rifting and strikeslip movements (from Howell 1989, ﬁg. 4.5). (With kind permission of Klumer Academic Publishers.)

Fig. 10.40 Hypothetical terrane assembly, showing provenancial linking, overlap sequence, pluton stitching, and welding metamorphism (from Howell 1989, ﬁg. 4.9). (With kind permission of Kluwer Academic Publishers.)

Tectonics 239

Fig. 10.41 Paleozoic terranes of Southeast Asia. Abbreviations: Kan, Kanchanaburi; CC, Cholburi; NWS, NES, SWS, northwest, northeast and southwest Sumatra, respectively; WB, West Borneo; LLP, Luang Prabang-Petchabun (from Brookﬁeld 1996, ﬁg. 5).

240 Chapter 10

Fig. 10.42 Reconstruction of Southeast Asian Palaeozoic terrane relationships in early Mesozoic times. (a) Removal of left-lateral motion on Bangka shear. (b) Removal of Oligocene to Miocene opening of the China Basin (Cenozoic rift basins are vertically hatched). (c) Removal of early Cenozoic basin rifting. (d) Removal of Cretaceous–early Cenozoic large left-lateral motion on Red River fault, and small right-lateral motion on Laityau Petchabun-Khiong Marui fault. (e) Removal of left-lateral motion on Mae Ping fault system. Note simpliﬁcation of sutures and alignment. (f) Possible latest Triassic reconstruction with Jurassic collision of northwest and southwest Sumatra with northeast Sumatra, etc., partly restored (from Brookﬁeld 1996, ﬁg. 6).

and paleomagnetism (Sengor et al. 1988). It is also important to work from the primary reports of ﬁeld geologists where possible, since many secondary sources carry so many interpretative assumptions as to be useless (Hamilton 1989). For example, southeast Asia is a particularly complex area due to the collision of three large plates at an unstable, ever-changing triple junction, in which a number of Paleozoic terranes are separated by major faults and

suture zones (Fig. 10.41). The terranes deﬁned here are simply noted as arc and oceanic, etc., although their individual stratigraphies were used to deﬁne them (Brookﬁeld 1996). Some idea of the early Mesozoic relationships of these terranes, which may help in Paleozoic analysis, may be obtained by progressively removing the Mezozoic–Cenozoic motions in order and reconstructing the terrane positions at different times (Fig. 10.42).

11 Sea-level changes

11.1 Eustatic or “absolute” changes of sea level 11.2 Relative changes of sea level

Sea-level changes affect the base levels of erosion, and hence sedimentation, in both continental and marine basins, and 71% of the earth’s surface is ocean. The shoreline marks the intersection of mean sea level and the land. Changes in mean sea level are changes in the oceanic geoid, and are EUSTATIC (ABSOLUTE) CHANGES OF SEA LEVEL. Changes in the land may involve simple isostatic adjustments or isostatically uncompensated local tectonic movements. It is the interaction of eustasy, isostasy, and local tectonics that drive the RELATIVE CHANGES OF SEA LEVEL with respect to the earth’s surface. Yet, in the past, these were studied separately by geophysicists, geologists, and geomorphologists with little understanding of, or communication among, one another (Fairbridge 1983). Ignoring isostasy and local tectonics, as some stratigraphy texts still do, creates problems in understanding relative sea-level changes. We cannot separate eustatic, isostatic, and local tectonics effects in rock sections: all we can identify are relative changes of sea level from baselevel changes, positions of shorelines, and facies relationships.

11.1 Eustatic or “absolute” changes of sea level Eduard Suess deﬁned eustasy in 1888 after observing synchroneity in Tertiary marine transgressions and regressions in the Paris Basin (for a history of the concept, see Dott 1992). Only the earth has oceans, so eustasy applies to no other planets – except perhaps to early Mars. Eustatic changes were previously believed to be worldwide and simultaneous, and to give rise to vertically displaced, but parallel, shorelines in the absence of tectonic deformation. Proving eustasy requires very good time control on widely separated successions, so it is difﬁcult to decide whether apparent sea-level changes are due to regional tectonics or are truly eustatic (Hallam 1992). Furthermore, a change in sea level automatically causes isostatic deformation, and satellite observations have shown that mean sea level (the oceanic geoid) is not ﬂat and uniform. It has relief and is always deforming or adjusting to gravitational and rotational changes (Fig. 11.1).

242 Chapter 11

Fig. 11.1 The difference between old and current concepts of eustasy (Morner 1983, ﬁg. 5.1). (Reprinted by permission of Oxford University Press.)

(a)

(b)

(c)

Fig. 11.2 Types of eustasy and their effect: (a) geoidal, (b) tectonic, (c) climatic (modiﬁed from Morner 1987, ﬁg.11.3). (Reproduced with permission of Blackwell Publishing Ltd.)

There are three types of eustasy: geoidal, tectonic, and climatic (Fig. 11.2). 11.1.1 Geoidal eustasy The geoid varies according to mass variations within the earth. Being an equipotential surface caused by the balance between attraction (gravity) and repulsion (rotation), the oceanic geoid responds immediately to any changes in these. It is worth noting at this stage the

importance of geoidal eustasy and its intimate and inseparable interactions with tectonic and climatic eustasy (Fig. 11.3). The formation of polar ice-caps not only extracts water and cools the oceans, causing a lowering of sea level (climatic eustasy), but it also transfers angular momentum from equatorial to polar regions (Wu & Peltier 1984). This requires the earth to spin faster and thus increase centrifugal forces so that a more squashed (oblate) spheroid is required to balance the forces (geoidal eustasy). Water immediately, and mantle more gradually, ﬂows towards the equator. An immediate rise in equatorial sea level during the onset of glaciation is followed by an uplift of the sea ﬂoor (and oceanic islands) and a drop in sea level as mantle replaces water (tectonic eustasy). Conversely, the melting of polar ice-caps transfers momentum from poles to equator. The earth spins more slowly, the spheroid becomes less oblate, and sea level at ﬁrst drops, then rises as water replaces mantle. On a smaller scale, gravitational attraction of land masses causes sea level to rise towards coasts, and high mountains with their roots and large ice-caps have greater effects (Fig. 11.2a). Gradual wearing down of mountains will cause sea levels to drop somewhat at the coast, as will a reduction in ice mass. However, the greatest effects are related to the large geoid anomalies, which are independent of the earth’s surface topography (see Section 10.1.3). Migration of these geoid anomalies can, by themselves, cause regional sea-level changes (Nunn 1986), which vary in magnitude depending on the hypsometry of the land and the size of the land mass

Sea-level changes 243

Fig. 11.3 Interaction of geoidal, tectonic, and climatic eustasy: (a) glacial; (b) interglacial (courtesy E. Grossman, Univeristy of Hawaii).

(Gurnis 1990). Two of the standard areas for determining Cenozoic eustatic sea-level changes are Barbados and New Guinea. Both areas have strong geoid anomalies (Barbados - 50 m, New Guinea + 70 m), so geoidal migration could explain some of the regional eustatic changes. Because of the interlocked effects of the three controls on eustatic sea level, worldwide sea-level changes of the same magnitude are a myth. Although changes of 200 m or more might swamp the maximum regional variations due to geoid migration (around 150 m), the use of supposedly uniform eustatic sea-level changes to calibrate stratigraphic successions worldwide, or even within one large continent, is both observationally and conceptually worthless, even without considering the imprecise dating on which it is based (Hubbard 1988). 11.1.2 Tectonic eustasy TECTONIC EUSTASY involves changing the capacity of oceans by: (i) increasing or decreasing their area; or (ii) increasing or decreasing their depths.

11.1.2.1 Areas of oceans Areas of oceans can be changed by compressing or expanding continents, or isolating small ocean basins, with resulting lateral mass transfers either by internal (plate tectonic) or external (surface) cycles. Compressing continents causes eustatic drops in sea level, while expanding continents causes eustatic rises in sea level. Thus, Africa is presently compressed by the surrounding mid-oceanic rises, and has a higher height/area ratio (hypsometry) than other continents. The resulting expansion in ocean area due to crustal and subcrustal squeezing and mass transfers into Africa should cause a eustatic drop in sea level. In contrast, Eastern Asia is gravitationally extending over the Pacific, and large rivers transfer mass from high areas to coastal deltas. The resulting contraction in ocean area should cause a eustatic rise in sea level. The apparent “worldwide” eustatic rise is actually the sum of these often-divergent effects. Closing off or ﬂooding of small ocean basins increases or decreases the total ocean area. Miocene tectonic closure of the Straits of Gibralter cut off the Mediterranean Sea. Subsequent complete evaporation of the ocean caused a eustatic rise in world sea level

244 Chapter 11

Fig. 11.4 Inferred ocean ridge length for (a) the Cretaceous and (b) the Permian (heavy black lines are ocean ridges) (courtesy R. Blakey, University of Arizona), and (c) cross-sectional areas for varying spreading rates (modiﬁed from Pitman 1979).

of about 12 m, while its subsequent ﬂooding reduced sea level by the same amount (Donovan & Jones 1979). 11.1.2.2 Depths of oceans The depths of oceans can be changed by isostatic changes due to loading/unloading and heating/ cooling. As Hallam (1963) noted, the quantitatively most important tectonic cause of eustatic sea-level changes is variations in the average depth of the sea ﬂoor, which primarily depends on the volume of the mid-oceanic rises (spreading rates and total length), and the mean age of the oceanic lithosphere (Fig. 11.4). Rapidly spreading ridges cause displacement of sea water and eustatic rises relative to slowly spreading ridges (Pitman 1979). The rapid spreading and resulting rapid plate movements during the Cretaceous

break-up of Gondwanaland may explain the great rise in sea level during that period. The stagnation in Permian plate motions (there are few Permian magmatic arcs) and apparent gap in spreading between Permian and Triassic times (with consequent subsidence of the aging sea ﬂoor and an increase in mean age) may partly explain the low sea level for much of this period (Hallam 1992). Ocean depths can also be reduced simply by adding lava or sediment to the sea ﬂoor. Vast outpourings of submarine lava formed extensive shallow-water plateaus in the southwest Paciﬁc during the Cretaceous, contributing to the Cretaceous sea-level rise. The amount of glacially transported sediments dumped at the continental edges during the Quaternary would lead to a rise in sea level in the absence of other factors. However, none of these mechanisms can change sea level at the rates inferred by Haq et al. (1988) for

Sea-level changes 245 (a)

Fig. 11.5 (a) Emergent ice-contact delta, Bafﬁn Island. (Reproduced with the permission of the Minister of Public Works and Government Services Canada, 2003 and courtesy of Natural Rosources of Canada, Geological Survey of Canada). (b) Dated driftwood shows shoreline positions during rebound (from Evans 1991, ﬁg. 3). (Reproduced with permission of Philip Allan Updates.)

non-glacial periods. Maximum rates of eustatic change based on observed spreading rate changes and their effects on oceanic ridge topography are about 10 m/ million years (Dewey & Pitman 1998) – an extremely slow rate in comparison with climatically driven eustasy (see Section 11.1.3). 11.1.3 Climatic eustasy CLIMATIC EUSTASY involves changing ocean volumes, either by: (i) simple thermal expansion/contraction of sea water; or (ii) adding/extracting water.

11.1.3.1 Thermal expansion or contraction of water Changing the average ocean temperature by 1 °C causes

volume changes that will raise or lower sea level by about 2 m independently of any other changes (Southam & Hay 1981). The same concepts can be applied (on a smaller scale) to lakes. 11.1.3.2 Adding or extracting water Adding or extracting water can occur in a number of ways. Over a geologically short time-frame, water is extracted from the oceans during glacial periods (glacial eustasy) and when large lakes exist in continental interiors (lacustrine eustasy). It is, of course, returned during interglacial and arid periods. Over a longer time-frame, additions from the interior of the earth have probably slowly increased the mass of the oceans, although subtractions can occur in porewaters and by hydration

246 Chapter 11 1939). If ancient hypsometries were similar, then the areas ﬂooded in the Ordovician and Cretaceous suggest maximum eustatic sea-level rises of about 100 m. However, hypsometries have probably changed, and eustatic sea-level changes can only be inferred from relative sealevel changes.

11.2 Relative changes of sea level Relative sea level is deﬁned by shorelines and may be independent of eustatic changes. If the land is subsiding, the relative sea level can be rising while the eustatic sea level is dropping, and vice versa (Pitman 1979). Relative sea-level changes can be studied using three main methods: the position of ancient shorelines; the successive areas of marine sediments; and the vertical changes of facies and inferred water depth in stratigraphic sections (Hallam 1992). 11.2.1 Ancient shorelines

Fig. 11.6 Forebulge migration and effects on relative sea level, northeast North America. Elevation (m) is relative to contemporary sea level at the site. Data points (squares and circles) are for Maine (from Barnhardt et al. 1995, ﬁg. 3). (Copyright © 1995. Reproduced with permission of the publisher, the Geological Society of America, Boulder, Colorado, USA.)

reactions with rocks (e.g. hydrothermal alteration in cooling oceanic lithosphere). Of these, only glacial eustasy signiﬁcantly affects stratigraphic sequences, as far as we know. Calculations based on the volume of oceans and of Quaternary glaciers give values for glacial eustatic sea-level changes of up to 200 m rises and falls between glacial and interglacial periods. Rates of change are geologically very fast: about 100 m per 20,000 years (0.5 cm/year) – this is fast! A eustatic rise from present sea level of 100 m would ﬂood 50% of the present land, based on modern hypsometry; and a 50 m rise would ﬂood 25% (Kuenen

The positions of ancient shorelines are available only for the Quaternary in adequate numbers for different times, where they provide fascinating examples of the interaction of varying rates of eustasy, isostasy, and local tectonics (Martini et al. 2001). In areas once covered by ice sheets, rapid postglacial eustatic rises of sea level were overtaken by larger amplitude though slower isostatic uplift. For example, in northern Canada, deltas that built out into the sea during the immediate postglacial sea-level rise, have now been raised high above present sea-level by isostatic rebound (Fig. 11.5). In southern Scotland, marine clays deposited on tree stumps (which grew in peat on glacial till) as the ice retreated, are now above sea level and covered by sand dunes and modern soils, and all are now being eroded as rebound continues. This history is even recorded in a medieval Scottish rhyme: “First a wood, then a sea. Now a moss and ever will be.” Postglacial migration and collapse of forebulges developed at the edge of the ice sheets caused diachronous relative sea-level changes from south to north. The initial marine transgression is followed, as the forebulge migrates, by a reduction in the rate of sea-level rise, or regression, depending on the relative rates of forebulge migration and sea-level rise, and then by renewed ﬂooding as the forebulge passes (Fig. 11.6). The same

Sea-level changes 247

Fig. 11.7 Inferred land and sea (stippled) areas for the Hettangian and Oxfordian (from Hallam 1992, ﬁg. 2.1). Note that these are not equal-area maps. (Reproduced with permission of the Edinburgh Geological Society.)

Fig. 11.8 (a) Earth hypsometry today and inferred for 220 Ma; (b) eustatic sea-level curve for the Ordovician (from Hallam 1992, ﬁgs 2.7, 3.4). (Reproduced with permission of the Edinburgh Geological Society.)

248 Chapter 11

Fig. 11.9 Cretaceous cycle with inferred paleodepths and sea-level changes (from Hallam 1992, ﬁg. 2.2). (Reproduced with permission of the Edinburgh Geological Society.)

effects are seen in front of thrust sheets and around large volcanoes.

sible to correlate adequately on a global scale (Fig. 11.8b).

11.2.2 Areas of marine sedimentation

11.2.3 Facies changes and water depths

Successive areas of marine sedimentation for speciﬁc time periods can be plotted on equal-area maps and the (reasonable) assumption made that a greater area indicates a higher sea level (Hallam 1992). Using Odin’s (1994) time-scale, the Lower Jurassic Hettangian Stage (3 Ma long) has a 5% marine coverage of the continents, while the Upper Jurassic Oxfordian Stage (8 Ma long) has a 25% marine coverage (Fig. 11.7). These estimates are approximate because erosion removes much sediment, particularly near shorelines, and extrapolations must be made beneath younger strata. Furthermore, changes in earth hypsometry through time affects estimates of sea-level rise (Fig. 11.8a), and adequate correlation of equivalent units is a problem. For example, Ordovician sea-level changes are based on poorly deﬁned stages, which are impos-

Vertical changes of facies and inferred water depth can be seen in individual sections (Fig. 11.9). The challenge is to correlate individual sections well enough to demonstrate that the change is actually eustatic and not caused by regional tectonics or sediment supply. In sequence stratigraphy, individual transgressive surfaces are considered practically isochronous and usable as time-lines. Unfortunately for this great idea, in welldated sections, the transgression can often be shown to be gradual, over millions of years (see Chapter 14 for discussion). Varying sediment supply (perhaps related to tectonic and/or climatic changes hundreds or even thousands of kilometers away) may cause relative changes of sea level, independent of eustatic changes. For example, the supposed site of Troy is too far from the sea to ﬁt the descriptions in Homer’s Iliad. However, when the Trojan

Sea-level changes 249

Fig. 11.10 Evolution of the area around Troy in the last 7000 years (from Roberts 1989, ﬁg. 6.3).

war occurred, around 7000 years ago, Troy was on an estuary – probably formed by rapid postglacial sealevel rise. Even when the story was ﬁnally written down around 3500 years ago, Troy was still coastal

and the story made sense to the contemporaries of Socrates and Plato. Since then, increased sedimentation (partly due to climatically induced, increased erosion rates) has completely ﬁlled the estuary with about 50 m

250 Chapter 11

Fig. 11.11 Effect of varying sediment supply on transgressions and regressions (from Vail et al. 1977, ﬁg. 3).

of sediment, causing a relative drop in sea level (Fig. 11.10). Relative changes of sea level interact with sediment supply in complex ways to alter base levels of erosion and facies distributions both on land and under water (Fig. 11.11). In turn, base-level changes cause sedimen-

tary packages to be separated by erosion surfaces or unconformities, and these can then be interpreted in terms of cyclical variations of sea level (Fig. 11.12). Such inferred cycles and supercycles (Fig. 11.13) are sometimes claimed to be worldwide, and have been used as a way of classifying and regulating the stratigraphic system

Fig. 11.12 Sequence stratigraphy in (a) space and (b) time (from Vail et al. 1977, ﬁg. 13).

Fig. 11.13 Regional cycles of sea level through time (from Vail et al. 1977, ﬁg. 13).

252 Chapter 11

Fig. 11.14 Worldwide sea-level chart for the Mesozoic–Cenozoic (from Haq et al. 1987, ﬁg. 5).

Sea-level changes 253

Fig. 11.15 Asymmetrical cycles due to erosion during regression (from Hallam 1992, ﬁg. 2.15). (Reproduced with permission of the Edinburgh Geological Society.)

(Fig. 11.14). The curious shape of the curves, with rapid transgression followed by gradual regression, is probably an artifact of the different interactions of eustasy with isostasy during transgressions and regressions. During transgression, an additional isostatic subsidence is added on to the sea-level rise, increasing the rate and amount of base-level rise, increasing the amount and rate of the relative sea-level rise, and decreasing the sedimentation rate so that condensed deposits accumulate. During regressions, an additional isostatic uplift is added to the sea-level drop, increasing the amount and rate of base-level drop, increasing the amount

and rate of sea-level fall, and increasing erosion and hence sedimentation in areas not yet reached by the regression. Most cycles are highly asymmetrical, except in deep water below the base-level changes (Fig. 11.15). Relative changes of sea level also affect organisms indirectly because of the effects on habitat. Thus, the main geologic systems, deﬁned on successive changes in faunas, are frequently bounded by unconformities due to relative drops in sea level. On a smaller scale, Jurassic and other system stages show changes in zone fossils across cycle boundaries (Fig. 11.16).

Fig. 11.16 Changes in Lower Jurassic ammonite genera through time (from Hallam 1961, table 2).

12 Climate

12.1 Present distribution and character of climates 12.2 Identiﬁable climatic effects on sediments, biotas, and stable isotopes 12.3 Controls on climate and climatic change CLIMATE is the long-term average combination of temperature, winds, and precipitation on land, with the addition of salinity, structure of water masses, and currents under water. Climate ﬂuctuates daily (night is usually colder than day) and annually (winter is colder than summer), and at various scales in between. It also ﬂuctuates over much longer time spans; for example, the ﬂuctuations recorded since the Mesozoic can span hundreds to millions of years (Frakes et al. 1992). We are currently in a glacial age of extreme polarization and variability of climate, but at other times, more equable conditions existed (Fig. 12.1). To understand climatic effects and controls, you need to know several things: 1 the present distribution and character of climates, and the effects of seasonality; 2 the way climate affects sediments, biotas, and geochemistry; and how these, in turn, can be used to identify climatic variables of the past; 3 astronomic, tectonic, eustatic, and biological controls on climate and how they operate, together with some examples of climatic control from stratigraphic successions.

However, there are many complicating factors which make interpretation difﬁcult. For example, the progressive build-up of polar ice-caps during ice ages creates steeper temperature gradients between the equator and poles, and more vigorous drying trade winds can mobilize sand into dunes in midlatitude deserts. It was only two to three million years ago that extremely arid conditions began in North Africa: the great tropical inland lakes of North Africa began to dry out at the same time as the start of glaciation to the north (Dawson 1992).

12.1 Present distribution and character of climates The basis of climate is how the atmosphere and oceans transfer energy around the globe. The atmosphere, when viewed from space, appears to consist of an endless succession of transient eddies swirling erratically across the face of the planet (Fig. 12.2). However, the dominant pattern is one of intense heat convection around the equator, which generates the Hadley

256 Chapter 12

Fig. 12.1 (a) Temperature for the last 180 Ma; (b) one million years; (c) 160,000 years (from Mackenzie & Mackenzie 1995, ﬁg. 10.3). (Reprinted by permission of Pearson Education, Inc., Upper Saddle River, NJ.)

Climate 257 circulation and migrates back and forth across the equator with the annual cycle of the seasons. Hot, moist air rises to form the cloudy equatorial belt, and then descends and dries polewards (around 30°) to form hot deserts. In the northern hemisphere (with an equivalent in the southern hemisphere), the northeasterly trade

Fig. 12.2 Surface temperatures and clouds on earth, October 29, 2003, (courtesy SSEC, University of Wisconsin-Madison. See http://www.ssec.wisc.edu/)

winds then ﬂow back towards the equator. In the summer, tropical cyclones develop in the easterly ﬂow and move northwards to merge with the mid-latitude westerlies, in which depressions carry air toward the poles to cool and sink, forming polar deserts (Fig. 12.3). Ocean patterns are harder to see, but consistent ones come from several centuries of observation (Fig. 12.4). These, like the atmosphere, are complicated by eddies generated by the mean ﬂows, and tides are important. However, in contrast to the atmosphere, sediments accumulate from shallow to deep areas in the ﬂuid, so the deeper ﬂow and structure also need to be considered. Temporary and permanent thermoclines separate water masses with different temperature regimes, and at present complicated deep ﬂows carry water from the North Atlantic into the North Paciﬁc (Fig. 12.5). At other times in the past, especially during non-glacial epochs, both atmospheric and oceanic structure and circulation patterns were quite different and the deep oceans may have been periodically warm and stagnant. The present distribution of climatic belts on land and sea is the result of a strong latitudinal heat variation (and consequent vigorous atmospheric and oceanic

Fig. 12.3 General circulation of the northern hemisphere (from Burroughs 2001, ﬁg. 3.2). (Reproduced with permission of Cambridge University Press.)

258 Chapter 12

Fig. 12.4 General circulation of the oceans (from Burroughs 2001, ﬁg. 3.3). (Reproduced with permission of Cambridge University Press.)

Fig. 12.5 Oceanic conveyor belt. Cold, salty water near Iceland, ﬂows at depth to the North Paciﬁc where it rises, mixes, heats, and ﬂows westwards (Frakes et al. 1992, ﬁg. 12.5).

Climate 259

Fig. 12.6 Major biomes of the world (from Simmons 1982, ﬁg. 3.1 with marine added). (Reproduced with permission of HarperCollins.)

mass transfers) and separated continents (and consequent more maritime land climates and smaller oceanic circulations) (Fig. 12.6). They cannot be used as a basis for identifying climates in the past or inferring controlling factors without taking into account differing continent and ocean distributions, areas and hypsometries, differing global heat transfers, and the different organisms present at any one time. For example, Cretaceous dinosaurs and Eocene conifers lived at high latitudes subject to perpetual winter night, but in climates signiﬁcantly warmer and wetter than those now operating in these high latitudes. Mean temperatures and seasonality are very difﬁcult to infer for pre-Cenozoic times. For example, there is some doubt as to whether some Paleozoic limestones with diverse fossil faunas were deposited in shallow warm water or in deeper colder

water, or in alternating conditions across tropical thermoclines. Furthermore, seasonality and events such as storms and tsunamis are not yet considered enough in paleoclimatic analysis. It is also a mistake to infer climates based on paleomagnetically determined paleolatitudes using Recent climatic distributions. Nevertheless, within one sedimentary basin, climate can sometimes be inferred from the sediment, fossil, and geochemical characteristics (Suttner et al. 1981).

12.2 Identiﬁable climatic effects on sediments, biotas, and stable isotopes This section brieﬂy summarizes some of the evidence on

260 Chapter 12

Fig. 12.7 Upper Permian paleogeography (270–260 Ma) (courtesy R. Blakey, University of North Arizona).

which climate can (and frequently cannot) be inferred for deposits of sedimentary basins (Parrish 1998). 12.2.1 Sediments Certain sediment types are often used as evidence of climate. 12.2.1.1 Evaporite deposits Evaporite deposits result from an excess of evaporation over precipitation in dry areas, and range from gypsum through halite to sylvite in marine environments. Thick desert dune sands form in the same areas. Most Recent evaporites and thick dune sands develop in the hot desert belts between 15° and 35°, but some occur at latitudes of up to 50° in the extreme aridity of central Asia, and also in Antarctica. Extreme aridity, like that of central Asia, probably characterized the interior of the ancient supercontinents regardless of latitude; for example, arid deserts occupied large areas of Gondwanaland in the late Paleozoic irrespective of latitude (Fig. 12.7). 12.2.1.2 Red beds Red beds with ferric oxide grain coatings and cements suggest oxidizing arid and semi-arid climates. However,

even those with pedogenic carbonates indicate only seasonal wetting and drying, and color may survive from the source weathering environments (Walker et al. 1978). 12.2.1.3 Paleosols Paleosols developed like modern soils under different climatic regimes and can provide a rough indication of climate (Fig. 12.8). However, soils are strongly controlled by vegetation and are therefore difﬁcult to use before the Cretaceous evolution of today’s dominantly ﬂowering plant vegetation. 12.2.1.4 Glacial tills and tillites Glacial tills and tillites are infallible indicators of ice and hence cold conditions. However, unless they are interbedded with marine deposits, they may simply indicate cold conditions at high altitudes. Only the stratigraphic correlation of widespread glacial units can establish widespread cold conditions. Otherwise they may simply record continental drift across polar regions or changing altitudes through time (Fig. 12.9). They are also easily confused with deposits of other agents such as mass ﬂows and impacts.

Climate 261

Fig. 12.8 Temperature and moisture effects on soil development during northern hemisphere deglaciation and glaciation (from Roberts 1989, ﬁg. 3.1).

Fig. 12.9 Migration of Paleozoic poles (really plate movements) across Gondwanaland with Ordovician and early Permian ice-caps shown (from Van Andel 1994, ﬁg. 7.3). (Reproduced with permission of Cambridge University Press.)

262 Chapter 12

Fig. 12.10 Nutrient-rich upwelling areas at present (from Waller 1996, ﬁg. 25).

12.2.1.5 Varved clays

12.2.2 Biotas

Varved clays form in seasonally frozen lakes. Fine sand and silt laminae are deposited in the deeper water of the open lakes by underﬂows associated with spring and summer freshwater input and wave action. These laminae are overlain by organic-rich clays deposited during winter as dead plankton and clay settles in the still water below the frozen surface. However, it is often difﬁcult to separate non-glacial turbidity current “varves” from true glacial ones.

Fossils are often used to infer climatic characteristics, but most have severe problems. Even total diversities are misleading. Antarctic benthic marine communities are as diverse and complex as tropical ones. The main control is stability not actual temperature.

12.2.1.6 Glauconite and phosphate Glauconite and phosphate form in areas of cold, nutrient-rich, upwelling ocean currents which are today concentrated off the western coasts of continents between 40°N and 40°S (Fig. 12.10). Plotting all these indicators on world paleotectonic and sea-level maps based on plate tectonics and/or paleomagnetism and areal distributions of marine facies can help in determining world climate. However, none of them allow actual values of temperatures to be determined.

12.2.2.1 Fossil plants Fossil plants can be used in the Cenozoic to infer climate based on angiosperm leaf anatomies (Fig. 12.11). However, similar climatic inferences for pre-Cenozoic ﬂoras are uncertain, except where other sedimentary features indicate climates for speciﬁc vegetation types. Peat and coal have been claimed to indicate warm and wet conditions. However, Recent peats range from the tropics to near the poles: some of the most extensive and widespread peat deposits are now forming in the Canadian and Russian arctic (Martini & Glooschenko 1985). Without detailed study of the plants, peat and coal only indicate wet, waterlogged conditions with rates of oxidation less than the rate of plant accumulation. The coal swamps of the Carboniferous were not necessarily hot (Cecil et al. 1985).

Climate 263 100

Entir-margined species (%)

30

Northern California

80

Pacific Northwest Southern Alaska

60

20

40 10 20

Estimated mean annual temperature (°C)

Mississippi embayment

0

Paleocene

Eocene

Oligocene

Miocene

Pilocene

0

Fig. 12.11 Estimated mean annual temperature from leaf anatomies for the Cenozoic of western North America from the Mississippi embayment (dots and solid line), the Paciﬁc Northwest (solid triangles and dotted line), northern California (open circles) and southern Alaska (open triangles) (from Wolfe 1978, ﬁg. 15.22).

12.2.2.2 Fossil animals Animals of various types and communities are used in similar ways to plants – with the same problems. Cenozoic molluscan faunas were used to infer winter sea temperatures at different times along the coast of western North America, and the climatic deterioration and gradient steepening is obvious (Fig. 12.12). However, pre-Cenozoic inferences become increasingly uncertain with age. Fossil corals (especially coral reefs) are frequently used to infer warm, shallow tropical conditions, based on the Recent distribution of warm, shallow hermatypic reefs. Even unrelated Paleozoic corals are used in this way, conveniently ignoring the extensive large Recent coral reefs of deeper and polar colder waters (Fig. 12.13). 12.2.3 Stable isotopes Stable isotopes exist in well-deﬁned ratios in the oceans and atmosphere, controlled by dynamic processes. These ratios change as the importance of the various processes changes through time. Ancient temperatures have been calculated from oxygen isotope variations preserved in carbon and carbonate.

Fig. 12.12 Cenozoic winter isotherms based on mollusca for western North America (from Goudie 1992, ﬁg. 1.8). (Reproduced by permission of Oxford University Press.)

12.2.3.1 Oxygen isotopes Oxygen isotopes consist of the lighter and dominant 16O (99.8%) and the heavier and rarer 18O (0.2%). During evaporation, water with the heavier oxygen isotope, H218O, is preferentially left behind, so rainwater and freshwater is enriched in H216O. There is normally a balance in oceans between the loss of H216O by evaporation and the gain of H216O supplied by rivers, and this dynamic balance is affected by temperature. Vigorous evaporation at warmer temperatures increases the contrast between ocean and freshwater. Carbonate deposited inorganically from sea water has an isotopic 18O/16O ratio close to that of the water from which it came (expressed as a deviation, d18O, from a standard). Organic carbon or carbonate may be close to or signiﬁcantly different from that of the water from which it came, depending on the physiology of the organism. If the isotopic ratio of the oceans is known for a standard temperature, then fossil carbon and carbonates can be used to calculate ocean temperatures. However, there are many problems with using oxygen isotopes for determining ancient ocean temperatures, even if the isotopic ratios are unaltered. First, ancient oceanic isotopic ratios have to be related to an assumed temperature – there is no practical way of determining them independently. Some ancient ocean temperatures turn out to be unbelievably high if modern ocean isotopic ratios are used. Second, if freshwater is locked up in continental glaciers (or lakes), then ocean water becomes relatively enriched in H218O regardless of temperature (Fig. 12.14). Quaternary oceanic isotopic variations are primarily controlled by the volume of ice locked up on land during glacial periods and not by

264 Chapter 12

Fig. 12.13 Distribution of coral reefs (open circles), deeper water coral banks (black dots) and solitary corals around Antarctica (triangles). The dashed line is the 21 °C annual isotherm (modiﬁed from Cairns & Stanley 1981, ﬁg. 2, and Garrison 1996, ﬁg. 18.17). (Reprinted with permission of Books/Cole, a division of Thomson Learning.)

Fig. 12.14 Oxygen isotope variation between (a) interglacial and (b) glacial phases.

Climate 265 temperature, and are used to deﬁne oxygen isotope stages which reﬂect ice volumes and, by inference, generally warm and cold periods. Third, the present oceans are stratiﬁed. Near-surface sea water, down to a variable level (or thermocline), roughly follows seasonal changes of temperature. Quaternary planktonic foraminifera thus show combined temperature and ice-volume effects. Deeper water, however, is colder and controlled by deep underﬂows of polar water. Benthonic foraminifera therefore have different (colder) isotopic ratios, which are independent of surface temperatures and reﬂect mostly ice-volume changes. These differences, developed during the Oligocene, are a good indicator of the onset of glacial conditions in Antarctica and the development of the present strongly thermally stratiﬁed ocean. In preOligocene times, these oceanic conditions need not apply: in fact, there is some evidence that oceans increased in both salinity and temperature with depth at times (Railsbach et al. 1990).

12.3 Controls on climate and climatic change Earth’s climate changes over time due to variations in total heat budget (which causes changes over the whole earth) and variations in heat budget distribution (which redistributes heat within the earth), either separately or in combination (Mackenzie & Mackenzie 1995). Four main factors affect the earth’s heat budget and its distribution: (i) changes in radiation received from the sun; (ii) changes in heat from the earth’s interior; (iii) changes in the proportion of “greenhouse” gases in the atmosphere; and (iv) changes in the distribution and hypsometry of land and sea. 12.3.1 Changes in radiation received from the sun Changes in radiation may be caused by either changes in the sun’s output, or changes in the radiation received on earth. Current theory suggests that the sun has heated up through time, its luminosity being about 30% greater today than in the Archean (which, paradoxically, has no good evidence of glaciation, perhaps because of greater heat ﬂow from the interior and an atmosphere richer in greenhouse gases). At present, shorter ﬂuctuations every decade reduce the output during increased sunspot activities. This affects earth’s climate, but observations have only been made

within the last few centuries, and the effects of sunspot activity cannot be determined from the geologic record. Intermediate-scale ﬂuctuations may likewise exist, and be the cause of longer term patterns (e.g. the Little Ice Age of the Middle Ages), but have not been recorded to date. Changes in radiation received can be caused by changes in space and atmospheric transparency to radiation. The solar system rotates around the center of the Milky Way galaxy about every 300 million years (the galactic year), passing through stationary nebulae of hydrogen-rich particles as it does so. This periodicity can be matched with major icehouse conditions back as far as 900 Ma; though even for this period it cannot account for the smaller glaciations such as those in the Ordovician, unless the cyclicity is 150 Ma. Large extraterrestrial impacts have changed climates at least temporarily by blocking radiation with dust thrown into the atmosphere, and the climatic change may have caused greater adverse effects at the end of the late Cretaceous than the immediate effects of the impact. Changes in the distribution of the heat received in different areas and at different times can affect climate. The earth’s orbit, obliquity, and inclination (MILANKOVITCH CYCLES) do not change the total heat budget, but they do change the distribution and timing of it over the earth’s surface in a regular and predictable way (Fig. 12.15) (Imbrie & Imbrie 1979). The various periodicities combine to give an insolation variation graph for a particular part of the world at a particular season. Fig. 12.15 is for the northern hemisphere summer; it is crucial to remember that the southern hemisphere has a very different pattern. In fact, one problem in Quaternary geology is that the ice ages of both hemispheres seem synchronized – which they should not be if the Milankovitch cycles were the primary control. Furthermore, northern hemisphere climates in the Quaternary oscillated between warm and cold periods with varying periodicities: a periodicity of 41,000 years prior to 900,000 years ago, after which the periodicity changed to 100,000 years, which is difﬁcult to explain. Control on sedimentary successions by Milankovitch cyclical variation is also periodically claimed for pre-Quaternary times (Fig. 12.16). 12.3.2 Changes in the earth’s heat ﬂow Heat ﬂow from the earth’s interior depends mostly on

266 Chapter 12

Fig. 12.15 Milankovitch effects on northern hemisphere summer insolation (modiﬁed from Covey 1984, p. 64).

Fig. 12.16 Possible orbital cyclicity in the Lower Cretaceous (Upper Albian) of Italy (based on Herbert & Fischer 1986, ﬁg. 1).

Climate 267

Fig. 12.17 O2 and CO2 variations in the Phanerozoic (from Van Andel 1994, ﬁg. 14.6).

variations in the activity and extent of volcanoes, which are controlled by the lengths of ridges and subduction zones, the rate of plate movements, and the activity of hot spots (Courtillot 1999). Times of reduced plate movements should mark cooling, while outpourings of vast ﬂood basalts above hot spots should accompany heating. There is some support for this. Both the late Proterozoic and Carboniferous– Permian ice ages followed the assembly of large supercontinents, and a reduction in plate boundary activity and extent. The warm Triassic followed the eruption of vast ﬂood basalts in Siberia. However, Quaternary glaciation followed a time of rapid plate movement and multiple plate boundaries in the Tertiary; and the enormous outpouring of the Ontong Java oceanic plateau basalts had limited effects on climate in the warm Cretaceous (Larson & Erba 1999). 12.3.3 Changes in “greenhouse” gases Changes in the proportion of “greenhouse” gases in the atmosphere are caused by volcanic activity, by photosynthesis, by deposition of limestones, and by weathering of various rocks (Martini et al. 2001). Volcanoes pump out carbon dioxide, so warm conditions should correspond with increased volcanic activity.

Photosynthesis concentrates the carbon from carbon dioxide into organic matter, which can be buried (as in peat and coal). The evolution of abundant multicellular algae precedes the late Proterozoic ice age, while the evolution of complex land plant communities precedes the Carboniferous–Permian ice age. However, there is no corresponding change during the Tertiary cooling, and furthermore no major organic deposits in the late Proterozoic. Deposition of limestones takes CO2 out of the atmosphere, so periods of signiﬁcant limestone deposition should form before and during cool periods. Shelf seas with thick limestone deposits precede the late Proterozoic, the Ordovician, the Carboniferous– Permian, and the Tertiary cooling (the latter in the Cretaceous). However, cooling in both the Tertiary and the Carboniferous seems to have been prolonged long after limestone deposition had declined. Furthermore, other times of extensive limestone deposition (e.g. in the Silurian and mid-Jurassic) occurred during warm phases. Fluctuations in atmospheric gases through time are partly dependent on organic cycles, and these feed back into controlling organisms. Thus, the development and proliferation of multicellular animals in the Cambrian may have been triggered by an oxygen increase due to cyanobacterial photosynthesis. The burial of vast amounts of organic carbon as

268 Chapter 12

Fig. 12.18 The Late Cretaceous world around 80 Ma (courtesy R. Blakey, University of North Arizona).

coal during the Carboniferous may have caused an increase in oxygen, restricting plants and animals to wetter areas – or they would spontaneously combust (Fig. 12.17). Weathering also changes the amount of CO2 in the atmosphere: weathering of siliciclastic rocks removes carbon dioxide, while weathering of carbonates and organic rocks such as peat and coal adds carbon dioxide. Rapid plate movements and continental collisions raise mountains and promote increased weathering. 12.3.4 Changes in the distribution of land and sea and hypsometries Changes in the distribution of land and sea and hypsometries affect oceanic and atmospheric circulation systems and climatic regimes. Plate tectonic effects on climate can be studied with models of the earth’s climate system. For example, three very simple models of a polar continent with an ice-cap, a polar continent without an ice-cap, and a tropical continent, give very different latitudinal temperature distributions (Hay 1996). Plate reconstructions based on oceanic magnetic anomalies and paleomagnetism are reasonably accurate for the last 200 million years, and can be used to test such models against the stratigraphic record (Fig. 12.18).

Vertical uplift and subsidence of large areas causes major changes of climate worldwide and not simply in the affected area. Reorganization of jet streams in the atmosphere, or of the large marine currents, may lead to drastic variations in climate in certain regions. Iceland has experienced many expansions and contractions of glaciers, even in historic times, due to shifting of the Gulf Stream. The uplift of mountains and plateaus not only raises large areas above snowlines but also changes atmospheric circulation. For example, the southeast monsoon develops over the Indian Ocean, Arabian Sea, and southern Asia in the northern hemisphere summer (Fig. 12.19). The winds originate in subtropical highs in the southern hemisphere, where the winds are under the inﬂuence of the Coriolis force. This force disappears at the equator and the southeast trade winds are steered northwards by the uplifts of the East African rifts into the northern hemisphere, where they are deﬂected northeastwards by the Coriolis force, picking up moisture from the Indian Ocean and dropping it over India. However, it is the interaction of many variables that determines what happens to climate. Eustatic sea-level changes affect climate by varying the supply of water for precipitation, and altering the seasonal extremes of climate. High sea level extends more equable, wetter conditions to continental interiors, regardless of the actual climate, by supplying more extensive areas for

Climate 269

Fig. 12.19 Relationship of the southeast monsoon circulation to uplifts (from Hay 1996, ﬁg. 17).

evaporation and by acting as enormous heat exchangers. Low sea levels increase continentality by reducing evaporation and heat exchange. Thus, the Cretaceous is much more equable than the Triassic (compare Figs 12.7 and 12.18). Major changes in oceanic circulation patterns have profound effects. Shutting off the warm, northward-ﬂowing Gulf Stream in the North Atlantic would drastically reduce the salinity of the northern seas, and turn off the cold, more saline North Atlantic Deep Water current (Fig. 12.5). Such changes in marine currents may explain the Younger Dryas cold episode from about 13,000 to 10,000 years BP, when tempera-

tures both plunged and rose very rapidly in Europe. Supercontinent assembly raises high and extensive mountain ranges in which faster weathering of siliciclastic igneous and metamorphic rocks reduces atmospheric CO2 and cools the earth. Continent breakup and dispersal leads to the formation of new oceanic basins, changes oceanic and atmospheric circulation, increases volcanism and mantle degassing, increasing CO2 and thus warming the earth. The end result is that climate depends on so many factors that it is difﬁcult to say which, if any, is the most important (Fig. 12.20).

270 Chapter 12 (a)

(b)

Fig. 12.20 Climate related to sea level (a, Vail et al. 1977; b, Hallam 1992), volcanism, and supercontinent assembly and break-up (from Hallam 1992, ﬁg. 6.8).

13 Biology

13.1 13.2 13.3 13.4

Atmosphere and ocean changes Bioclastic sediment changes through time Sediment mixing by organisms Biogeographic changes

Organisms control stratigraphic successions in four main ways: (i) by altering the composition of the atmosphere and oceans; (ii) by forming different sediments through time due to evolution; (iii) by increasingly mixing sediments through time; and (iv) by forming different biotas in different areas through divergent evolution. This chapter gives a few brief examples to show that biological changes through time must be considered not only in biostratigraphy, but also in interpreting ancient environments and the resulting stratigraphic successions, in particular, those of the Proterozoic and Archean (Chapter 14).

13.1 Atmosphere and ocean changes The activities of organisms have, through time, completely altered the composition of the atmosphere and hence the dissolved gases of seas, lakes, rivers, and groundwater. This affects all surface processes, including weathering, erosion, transportation, and sedimentation. Throughout the Archean, cyanobacteria

progressively exchanged carbon dioxide with oxygen until the atmosphere and oceans contained free oxygen by the early Proterozoic (Fig. 13.1). Archean sediments contain detrital grains such as uraninite and metal sulphides that are oxidized in postArchean weathering proﬁles. Banded ironstones are formed mostly by the late Archean transitional phase to an oxygenating atmosphere; their banding being caused by oxidation of dissolved ferrous iron during daytime photosynthesis, alternating with silica precipitation during acid conditions at night (Harder 1995). The Silurian rise of land plants led not only to modiﬁcation of weathering processes, but also to an organic nutrient ﬂux to nearshore marine environments which altered their ecology. The rise in atmospheric oxygen during the Carboniferous increased weathering rates and may account for the very mature quartzites found with its coals.

13.2 Bioclastic sediment changes through time Bioclastic particle composition, structure, size, and shape

272 Chapter 13

Fig. 13.1 Earth atmospheric evolution, range of detrital uraninite and pyrite, end of banded iron formation, and comparison with inferred Martian atmospheric evolution (from Kirchvink & Weiss 2001, ﬁg. 2). (Reproduced with permission of the author.)

are initially controlled by the type of organisms. Abundant land vegetation did not exist until the Devonian, so coals form only from then onwards. Vegetation was not available to stabilize riverbanks until then, so most pre-Devonian river deposits are those of braided streams. Almost all limestones form by accumulation of carbonate skeletons or fragments derived from them. So, their composition and grain size is controlled in the ﬁrst instance by the nature of the organisms that form them. Organisms, and the environments they inhabit, have changed through time. Precambrian limestones are dominated by large dome-shaped stromatolites secreted by photosynthetic cyanobacteria, and the blocky particles and dust derived from them (Fig. 13.2). Obviously, such organisms (if photosynthetic) are restricted to the photic zone, so there are no deeper water, in-situ Precambrian limestones. Palaeozoic limestones have a wider distribution, but are dominated by marine, shelled organisms that secreted stable calcite, so cementation and early diagenesis tended to be slow and simple (James 1997). Reef corals became extinct at the end of the Devonian, and did not recover until the Permian, so there are no Carboniferous successions with true coral reefs. Abundant carbonate-secreting plankton did not appear until the mid-Mesozoic, so there are no chalks or equivalent deep-water limestones in the Paleozoic. Post-Paleozoic limestones are dominated by marine, shelled organisms that secreted metastable

aragonite, which dissolves easily to provide carbonaterich porewaters: early cementation and diagenesis is rapid and complicated. As Ager (1973) noted, certain geologic periods tend to be dominated by certain lithologies – almost all the result of control by organisms (Fig. 3.45).

13.3 Sediment mixing by organisms Organisms mix and churn sediments and destroy bedding and other structures, resulting in sediments that may be completely altered by the activities of organisms (trace fossils where identiﬁable) (Bromley 1990). The depth of mixing by organisms increases progressively through the Paleozoic and exponentially in the Triassic after the end-Permian extinctions: almost all ﬁne lithologies are heavily and deeply bioturbated from Mesozoic times onwards (Fig. 13.3). Only in the Precambrian can ﬁne tidal rhythmites remain so undisturbed by burrowing organisms that they can be used to infer tidal periodicities, lunar cycles, and year length (Williams 1998).

13.4 Biogeographic changes The independent evolution of organisms in separated areas forms biogeographically distinct biotas that can be

Biology 273 (a)

(b)

(c)

Fig. 13.2 (a) Modern stromatolites, Shark Bay (courtesy Hamelin Pool Marine Nature Reserve, Western Australia). (b) Cyanobacterial strand, Anabaena sp. ¥ 1000 (courtesy Cyanosite). (c) Stromatolite, North Pole, Australia, 3.5 Ba (courtesy T. Patterson, Carleton University, Ottawa).

Fig. 13.3 Increase in burrowing depth through time (from Ausich & Bottjer 1991, ﬁg. 2).

274 Chapter 13

Fig. 13.4 Paciﬁc and Atlantic trilobite faunal provinces in the late Cambrian Iapetus (modiﬁed from Windley 1984, ﬁg. 13.2). (© John Wiley & Sons Limited. Reproduced with permission.)

Fig. 13.5 Dynamic interpretation of changing latitudes of Tibetan terranes relative to the southern edge of paleoAsia (Angara terrane) (from Dewey et al. 1988, ﬁg. 3). (Reproduced with permission of The Royal Society.)

Biology 275 used to recognize once separated, now juxtaposed areas, and vice versa (Smith 1988). The opening and closing of a Paleozoic Atlantic was ﬁrst suggested by the very different Cambrian faunas across Appalachian– Caledonian structural belts, which also deﬁned terranes left on the wrong side of the present Atlantic when it opened (Fig. 13.4). Such biogeographic information can also assist in reconstructing the movement of crustal

blocks. For example, the various continental terranes of Tibet separated from the southern edge of the ancient Tethys at different times from the Carboniferous onwards. Their separation, northward ﬂight, and amalgamation with Asia can be inferred from a combination of paleomagnetism, faunal divergence, independent evolution, and convergence of biogeographic provinces among the various blocks (Fig. 13.5).

14 Stratigraphic problem times and places

14.1 14.2 14.3 14.4

Quaternary stratigraphy Archeological stratigraphy Proterozoic stratigraphy Archean stratigraphy

For most of the Phanerozoic, we use various stratigraphic methods to reconstruct histories to within several million years on average – which seems vaguely satisfactory. At other times, these methods give unsatisfactory answers. For the later Tertiary and particularly the Quaternary (about 2 Ma to present), methods of obtaining time relationships and ages are inadequate to separate events at the precision required, except for those within the last 35,000 years or so. Archeological stratigraphy works on even ﬁner spatial and temporal scales. For the Proterozoic, the lack of a usable biostratigraphy means that chronostratigraphy has to be based on numerical dating, except for the latest Proterozoic, where a gross biostratigraphy based on multicellular biotas is possible. In the Archean, additional problems are that igneous, metamorphic, and tectonic processes differed from modern ones because of greater heat ﬂow and more vigorous convection. Surface processes differed because of an anoxic atmosphere and the small size of continental crustal blocks, producing various strange lithologies, facies relationships, and basins.

14.1 Quaternary stratigraphy The Quaternary (1.65–0 Ma) consists of the glacial and interglacial deposits of the northern hemisphere ice age and its correlative deposits. The basic stratigraphy of glacial deposits is hard to work out. Deﬁning suitable rock units is difﬁcult because rapid lateral and vertical changes of environment (caused by very different juxtaposed processes and relatively fast eustatic, isostatic, and climatic changes) produce lensitic, erosionally bounded landforms and units. To compound this, glacial deposits rarely have distinct marker beds, and dating is difﬁcult, not only because facies relationships are complex, but because few dating methods are useful for glacial deposits. The only really useful deposits for space and time correlation, and dating, are interbedded and laterally equivalent non-glacial deposits. For example, distinctive condensed chemical and biochemical sediments deposited on sediment-starved shelves during interglacial rises of sea level are often the most useful for regional correlation. The complexity of glacial stratigraphy also means that

Stratigraphic problem times and places 277

Fig. 14.1 Quaternary section at Scarborough Bluffs, Ontario, Canada (modiﬁed from Eyles & Westgate 1987, ﬁg. 1). (Copyright © 1987. Reproduced with permission of the publisher, the Geological Society of America, Boulder, Colorado, USA.)

the rigid application of stratigraphic codes has never been popular (Eyles & Eyles 1992). The lack of good dating combined with complex lensitic and erosion-bounded units means that interpretations of successions can be wildly divergent. For example, take Scarborough Bluffs, which is a reference section for the early to mid-Wisconsin glacial deposits of North America (Fig. 14.1): one view explains the vertical sequence of tills alternating with river, delta, and lake sediments as being caused by regional advances and retreats of grounded ice sheets; another view explains the sequence as caused simply by local glacier ﬂuctuations within one lake. These sections are the standard for late Quaternary history across the North American mid-continent, so the controversy about subdivision is important. The original chronostratigraphic units of the Quaternary were climatic and based on biotas of continental

deposits. Thus, GLACIAL STAGES were deﬁned by major and prolonged expansion of glaciers; while the intervening INTERGLACIAL STAGES were deﬁned by conditions as warm as the present conditions (cooler and warmer glacial intervals are called STADIALS and INTERSTADIALS respectively). However, these are diachronous and vary inconsistently in various parts of the world, and they cannot be applied in any objective way to submarine strata. Few organisms evolved rapidly enough to erect useful Quaternary biostratigraphic units. Most vertical changes in biotas in any one area were caused by organisms “tracking” migrating environments, not by evolutionary changes. On land, the most useful biostratigraphy comes from mammals that evolved rapidly and spread quickly, such as small rodents whose teeth are frequently preserved, often in owl droppings. In Eurasia, the water vole has especially striking transitions at about 3.25 Ma and between 800,000 to

278 Chapter 14

Fig. 14.2 Quaternary isotope stages and land climatic units for the last 130,000 years (from Lowe & Walker 1997, ﬁg. 7.2). © Addison Wesley Longman 1985, 1997, reprinted by permission of Pearson Education Limited.

600,000 years ago, which are useful for European correlation. In Africa, elephant and pig lineages can be used to date sediments within about 500,000 years (Lister 1992). In the sea, the microfaunal-based units are little better than the land ones and almost impossible to cross-correlate. A better chronostratigraphic standard is based on oxygen isotope variations in relatively complete oceanic sediments: these deﬁne oxygen isotope stages, which reﬂect ice volumes and, by inference, warm and cold periods (Shackleton 1967). The stages are numbered backwards, with the warm stages odd-numbered and the cold stages even-numbered. The scheme now extends back to isotopic stage 102 at the Gauss–Matyuma boundary (at 2.58 Ma). However, this scheme cannot be directly correlated with the land-based scheme, because dating methods are not precise enough (Fig. 14.2). Correlation depends on identifying the individual climatic and eustatic sea-level changes inferred from the isotopic record. But Quaternary rates of sea-level change are so fast that they are not resolvable by the available dating methods, apart from carbon isotope dating within the last 35,000 years. Rapid melting of continental ice sheets caused sea levels to rise at an estimated rate of

about 1 cm/year, which is rapid enough to ﬂood most continental shelves within 10,000 years, and is thus almost instantaneous geologically. Although these are basically “event” horizons, we cannot resolve individual transgressions prior to the last one. Sidereal methods of dating, tree ring and varve chronologies, are limited to about the last 10,000 years. Geochronology is especially difﬁcult in the Quaternary because of the precision required. Only carbon isotope dating is precise enough to separate signiﬁcant events. However, the half-life of 14C is 5730 ± 40 years, so the method is reasonably accurate only within the last 35,000 years or so (i.e. six half-lives). Also, carbon isotope dating does not give calendar years: checking 14C dates against tree-ring ages shows systematic errors caused by changing 14C/12C atmospheric ratios over time. Uncalibrated 14C ages give large systematic errors, and remain a problem before 7000 BP. There are no reliable and accurate Quaternary dating methods beyond the limit of radiocarbon dating. All are too imprecise to separate intervals of less than several thousand years. Another approach is astrochronologic calibration. ASTROCHRONOLOGY is the calibration of geologic events to

Stratigraphic problem times and places 279

Fig. 14.3 Quaternary time-scale with astrochronological and K/Ar calibration (from Lowe & Walker 1997, table 1.1). © Addison Wesley Longman 1985, 1997, reprinted by permission of Pearson Education Limited.

computed astronomic time series; so far based only on past variations in the earth’s orbital parameters (Milankovitch cycles). In the Quaternary, Milankovitch cycles can be matched against ice volumes as recorded in the marine oxygen isotope stages. Some of these stages correlate with paleomagnetic reversals, and the basalts that carry these can be dated with reasonably precise

40Ar/39Ar

radiometric methods. Astronomic and radiometric dates agree well as far back as the Miocene (Hilgren et al. 1997). Nevertheless, argon dating still has standard uncertainties of up to 1.5%, despite decreasing analytical errors, even without noting geologic errors. Quaternary time-scales are still liable to revision (Fig. 14.3).

280 Chapter 14

Fig. 14.4 Cross-section of a limestone cave showing extreme lateral and vertical variation in horizons – and this is only a two-dimensional section (from Rapp & Hill 1998, ﬁg. 3.7).

Fig. 14.5 Meticulously described cellar cross-section in the Roman fort at Segontium, Wales, with coins calibrated. Note the lateral inﬁlling by earth from the doorway prior to building collapse (after ad 290, according to the youngest coin in a hoard in the earth below) (from Wheeler 1923, ﬁg. 17).

Stratigraphic problem times and places 281

Fig. 14.6 Proterozoic stratigraphy (from Levin 1991, ﬁg. 8.33).

14.2 Archeological stratigraphy Archeological stratigraphy is no different from any other stratigraphy, except that it requires greater accuracy, precision, and resolution, and can often be calibrated and integrated with historic records (Stein 1993). However, for pre-literate times, methods of dating are usually too imprecise to resolve important cultural changes since the tree-ring calibration of radiocarbon dating extends back to only 7000 years BP (Blackwell & Schwarcz 1993). Archeological excavation also destroys the record, which is rarely much of a problem in geologic stratigraphy except in unusual circumstances. This, plus the greater detail required, means that section excavation and description must be very thorough. GROUND-PENETRATING RADAR is now routinely used to locate promising areas and plan excavation prior to digging (Conyers & Goodman 1997). The techniques

are basically the same as those of seismic stratigraphy, and have the same problems. The high-frequency wavelengths used (generally 10–1000 MHz) give more precise information, although for that reason are limited to the shallow subsurface (usually less than 30 m depth). Ground-penetrating radar is also used for high-resolution near-surface Quaternary studies (Davis & Annan 1989). As in seismic stratigraphy, radar vertical and horizontal cross-sections, as well as three-dimensional radar reconstructions, are used to identify interesting areas (Plate 14.1). Archeological sections are often through heterogenous deposits with very rapid facies changes, cut by many numerous non-depositional and erosional diastems and unconformities (Piggott 1959). Particularly problematic are the deep pits and trenches cut by humans for various purposes, and which require careful separation from adjacent strata. Lithological correlation between even adjacent sections can be

282 Chapter 14

(b)

Fig. 14.7 (a) Archean rock distribution (from Goodwin 1976, ﬁg. 7). (b) Pilbara craton (3.5 Ba) from space; large granite domes with intervening greenstone belts (courtesy NASA). (Copyright 1976 © John Wiley & Sons Limited. Reproduced with permission.)

very difﬁcult, and until recently, few archaeologists studied sedimentology and did not realize that lateral migration, not vertical accretion, is the normal process of sedimentation. Many cave sites show both (Fig. 14.4). Archeological dating has the same problems as

dating in geology. Some artifacts (such as stone sickles and querns) deﬁne biofacies, since they record speciﬁc environments (here cereal cultivation). Others, such as coins and pottery, can deﬁne relative time units (chronostratigraphic units) within one cultural “basin”. However, these have the same limitations as

Stratigraphic problem times and places 283

Fig. 14.8 Archean facies models and stratigraphy. Speculative environments and relationships in the Western Churchill Province, Canada (from Lewry et al. 1985, ﬁg. 8). (Reproduced with permission of the Geological Association of Canada.)

zone fossils in geology, as they can be reworked (particularly coins, which resemble conodonts in durability): a layer can basically be no older than the youngest durable artifact in it. Coins can often be calibrated to a numeric (geochronologic) scale based on historic records. For example, Roman coins can be calibrated against abundant historic records (Fig. 14.5). With meager historic records, even the ﬁnest coin series is of little use (e.g. the magniﬁcent coin series of the Greek Bactrian kings of Central Asia).

14.3 Proterozoic stratigraphy The Proterozoic era (2500 Ma to the base of the Cambrian) occupies almost half of earth history, but it is the poor relation of the Archean era below (different

and exciting) and the Paleozoic era above (well known and familiar). The Paleoproterozoic (2500–1600 Ma) marks the transformation of the atmosphere and the ﬁnal assembly of large continents, with banded iron formations and the oldest known glacial deposits. The Mesoproterozoic (1600–900 Ma) and Neoproterozoic (900–540 Ma) resemble the Phanerozoic, in that the basic physical processes of plate tectonics and sedimentation appear to be similar, but with some differences. Supercontinents had more extensive blankets of mature sediments than in later eras and very unusual events, such as global ice ages (“snowball” earth), may have occurred. The Neoproterozoic marks the rapid biologic diversiﬁcation from colonial algae (about 1000 Ma) to complex skeletalized organisms. Low-resolution biostratigraphy is possible with a succession of soft-bodied fossils and succeeding Precambrian shelly fossils from

284 Chapter 14 about 700 to 540 Ma, but their ranges are still poorly known. The importance of biostratigraphy is painfully obvious from the use of “about” so often above. Without biostratigraphy, only expensive and time-consuming methods of relative dating, such as chemostratigraphy and magnetostratigraphy, are possible, and these require rough estimates of age before they can be used with any accuracy. Such age estimates need radiometric dating methods, which have variable accuracy and signiﬁcance (Chapter 7). The consequence is that Proterozoic divisions and units have to be mainly geochronologic and long (Fig. 14.6).

14.4 Archean stratigraphy “It is the business of the Archean geologist to remain skeptical of the over-rigid application of uniformity.” (Nisbet (1987))

The Archean is basically everything older then the base of the Proterozoic (before 2500 Ma), but there is very little left of it (Fig. 14.7). The problems of Archean stratigraphy are similar to those of the Proterozoic but compounded by the dominance of metamorphosed and deformed rocks. Although almost all the preserved Archean record post-dates the large impact phase of early solar system evolution (ending at 3800 Ma), the Archean meteorite ﬂux was undoubtedly greater than at younger times (Melosh 1990). Greater heat ﬂow, vigorous mantle convection, minute continents, and a reducing atmosphere meant that plate tectonic and sedimentary processes were very different (Condie 1997).

Stratigraphy in the Archean, even more than in the Proterozoic, is almost entirely chronostratigraphic since horizons can rarely be correlated to within 5–10 Ma. Facies variations within such divisions are based on the use of facies models (Fig. 14.8). Yet these must be applied to unusual greenstone belts with komatiitic basalts and bimodal volcanics, and strange juxtapositions and relationships of deep and shallow facies (Hyde 1980). Archean greenstone belt stratigraphies change through time (Lowe 1982). Older belts (3500– 3300 Ma) have volcaniclastic, cherty, and evaporitic sediments with local stromatolites, formed under shallow-water conditions on ﬂat basaltic volcanic platforms. Younger belts (2800–2600 Ma) have volcaniclastic sediments with banded ironstones and massive sulphides formed in deeper water around active sialic platforms of moderate relief, and with few shallow-water stromatolites and evaporites. These two types are separated by intrusive and metamorphic episodes of crustal accretion. Plate tectonic interpretations of such Archean greenstone belts have been tried, but with limited success (Erikkson et al. 1994), and may be inadequate (Hamilton 1998). Much of Archean stratigraphy remains regional in scope and many areas have not been investigated at all. The Archean stratigrapher, like the extraterrestrial stratigrapher (see Chapter 15), has the chance of a truly unknown area to study and of discovering new types of succession, and novel histories, unlike most stratigraphers who are stuck with redeﬁning rock units, reﬁning age relationships, or revising ages. Both Archean and extraterrestrial areas are now among the most exciting and rewarding ﬁelds of study in stratigraphy.

15 Extraterrestrial stratigraphy

15.1 15.2 15.3 15.4

The Moon Mercury Mars Venus

Within the last 30 years, spacecraft have passed and photographed almost all planets and moons within the solar system, and landed on the Moon, Mars, and Venus. The resulting cascade of remotely sensed data has revolutionized our knowledge of these bodies and started the new ﬁeld of extraterrestrial stratigraphy. This is not only important in itself, but has implications for deciphering earth history, especially in recognizing the importance of impacts and other processes in determining sedimentary basin development in both Archean and later times (Ernst et al. 2001). Curiously, extraterrestrial stratigraphy barely registers in current sedimentology or stratigraphy texts, although it forms an important section in many books on planetary geology (e.g. Christiansen & Hamblin 1995; Pater & Lissauer 2001). Extraterrestrial stratigraphy, in fact, recapitulates the challenge and excitement of early 19th century stratigraphy when the main eras and periods of earth history were established. This chapter attempts to convey that challenge and excitement with the inner planets and the Moon, although the same principles apply to the surfaces of the outer solar system moons (Fig. 15.1).

The surfaces of the inner planets and the Moon become ﬁxed at progressively younger stages due to deareasing rates of heat loss with size, yet they all show very different evolution (Tozer 1985; Lowman 1989; Rampino & Caldeira 1994) (Fig.15.2). Their stratigraphy is almost entirely based on sequences of events derived using Steno’s principles, and with relative ages of surfaces based on impact crater densities (with allowances for planet size and atmosphere). Impact crater densities are related to the decrease in meteorite (planetismal) bombardment through time shown by all solid bodies, calibrated entirely with the few radiometric dates from the Moon (Fig. 15.3).

15.1 The Moon The Moon (diameter 3475 km) has two major surface types: the bright highlands which form over 80% of its surface, and the darker seas (or maria) which form less than 20% of its surface. The highlands are saturated by meteorite craters and consist of anorthosites and

286 Chapter 15

Fig. 15.1 The diverse surfaces of Mars, Venus, and Earth at the same scale (courtesy NASA).

derived breccias formed by meteorite pulverization over 3900 Ma ago (Fig. 15.4a). The maria are concentrated on the earth-facing hemisphere and consist of ﬂood basalts mostly between 3900 and 3100 Ma in age, which inﬁll enormous impact basins (Fig. 15.4b). The maria are smooth because they have relatively few impact craters, so the main meteorite ﬂux must have been over by about 3900 Ma. The stratigraphy of the Moon is relatively simple, and radiometric dating allows calibration of the periods, which should probably be called eras considering their duration (Fig. 15.4c) (Wilhelms 1987). Fig. 15.2 Ages of inner planet and Moon rocks. The surface of Venus underwent catastrophic resurfacing around 500 Ma.

Extraterrestrial stratigraphy 287

Fig. 15.3 The Moon’s inferred meteorite ﬂux through time (data from Christiansen & Hamblin 1995, ﬁg. 4.25).

15.2 Mercury Mercury (diameter 4480 km) looks like the Moon, but without the maria/highland contrast. The largest feature is the 1300 km diameter Caloris impact basin. Focused shock waves from this impact probably formed the antipodal chaotic area of rock blocks and hills (Fig. 15.5). A similar process on earth may have triggered the Deccan ﬂood basalt eruptions of India, antipodal (after plate restoration) to the end Cretaceous Yucatan impact. Craters on Mercury are shallower than on the Moon, and secondary impact craters and rays are closer to the primary craters because of the increased gravity. Mercury’s heavily cratered terrain is interspersed with smooth plains, but no lava ﬂows have been seen. Unique faults scarps form linear features hundreds of kilometers long, and from a few hundred meters to a few kilometers in height. Mercury’s high uncompressed density may be due to a giant early impact stripping away much of its mantle. Since only one early spacecraft visited and photographed Mercury, its stratigraphy is somewhat uncertain. However, like the Moon, its main features formed very early and have not changed much since (Neukum et al. 2001).

15.3 Mars Mars (diameter 6370 km) is almost twice the size of Mercury and has a thin atmosphere. A number of spacecraft have visited and photographed the surface through its usually clear atmosphere, and several landers have analyzed its surface. So more is known about Mars than

any other planet except earth (Tanaka 1986; Carr et al. 1993; Albee 2001). Mars has two very different terrains: an older, cratered, southern hemisphere, like the Moon, and a younger, lower, northern hemisphere with igneous and tectonic features such as huge volcanoes and rift valleys (Fig. 15.6). The volcanoes seem to be several hundred million years old at least, and the rifts show large mass ﬂows and, like the polar regions, stratiﬁcation in valley sides (Fig. 15.7a). Both hemispheres, but especially the southern one, show evidence that enormous water ﬂoods occurred early on in its history (Fig. 15.7b) and there is disputable evidence of glaciers and small oceans (Carr 1996; Zent 1996). That water exists frozen beneath the surface is apparent from mass slurries and ﬂoods extending out from some craters. The stratigraphy of Mars is complex, of great interest, and continually being changed (Fig. 15.8). Oh, that I could live long enough to go there. Some of you may.

15.4 Venus Venus (diameter 12,103 km) is earth’s twin (diameter 12,742 km) but is a hellish place beneath its thick cloud deck, with surface temperatures above 300 °C (Fig. 15.9). Radar mapping at high resolution (100 m) reveals that most of the surface is within a kilometer of the mean radius, but large continental areas (8% of the surface) rise well above (3–5 km) the average surface level, and the maximum elevation contrast of around 13 km is about the same as on earth. The surface is dominated by complex plains (probably of volcanic origin) and there are shield volcanoes, and other possible volcanic

(b) (a)

Fig. 15.4 (a) Highlands, with Orientale basin; (b) Maria Imbrium towards Copernicus; (c) stratigraphy of the Moon.

Extraterrestrial stratigraphy 289 (a)

Fig. 15.5 (a) Cratered surface of Mercury with Caloris basin; (b) blocky and hilly terrain opposite Caloris; (c) origin of blocky terrain.

(b)

290 Chapter 15

Fig. 15.6 Mars: Vallis Marineris rift (as long as the USA is wide) across center; Tharsis volcanoes on left; part of cratered hemisphere on right (courtesy NASA).

(a)

Fig. 15.9 The surface of Venus, via radar images (courtesy NASA).

(b)

Fig. 15.7 (a) Mass ﬂows and stratiﬁcation on the sides of Valles marineris; (b) branching networks indicating slow erosion by rivers, Thaumasia area (courtesy NASA).

Extraterrestrial stratigraphy 291

Fig. 15.8 Uncertain Mars stratigraphy compared to that of the Moon (from Tanaka 1986, ﬁg. 2). (© (1986) American Geophysical Union. Modiﬁed by permission of the American Geophysical Union.)

features such as coronae (large circular structures with multiple rings), as well as long linear mountain ridges (Nimmo & McKenzie 1998). However, the random pattern of meteorite impacts on the surface of Venus suggests that most of the surface is of the same age. From the density of the patterns, a global resurfacing by lavas over the entire planet must have taken place around 300–600 Ma. Venus thus has almost no preserved geologic history, and thus no long stratigraphy (Basilevsky & Head 1995).

The moons of the outer planets have equally weird and interesting geology (e.g. Europa: Prockter & Pappalardo 2000). Extraterrestrial stratigraphy helps in understanding earth stratigraphy, particularly in interpreting the Archean, de-emphasizing strict uniformitarianism, and considering the inﬂuence of external processes such as meteorite impacts. It is now impossible to interpret the stratigraphy of the earth without looking outside it (Beatty et al. 1999).

Appendix 1 Imperial/metric conversions Units 1 kilometer (km) = 1000 meters (m) 1 meter (m) = 100 centimeters (cm) 1 centimeter (cm) = 0.39 inch (in) 1 mile (mi) = 5280 feet (ft) 1 foot (ft) = 12 inches (in) 1 inch (in) = 2.54 centimeters (cm) 1 square mile (mi2) = 640 acres (a) 1 kilogram (kg) = 1000 grams (g) 1 pound (lb) = 16 ounces (oz) 1 fathom = 6 feet (ft)

Conversions When you want to convert

multiply by

to ﬁnd

Length inches centimeters feet meters yards meters miles kilometers

2.54 0.39 0.30 3.28 0.91 1.09 1.61 0.62

centimeters inches meters feet meters yards kilometers miles

Area square inches

6.45

square centimeters

0.15

square centimeters square inches

square feet

0.09

square meters square feet square kilometers square miles

square meters square miles

10.76 2.59

square kilometers

0.39

Volume cubic inches

16.38

cubic centimeters

0.06

cubic feet

0.028

cubic meters cubic miles

35.3 4.17

cubic kilometers liters liters US gallons

0.24 1.06 0.26 3.78

cubic centimeters cubic inches cubic meters cubic feet cubic kilometers cubic miles quarts gallons liters

Mass and weight ounces grams pounds kilograms

28.35 0.035 0.45 2.205

grams ounces kilograms pounds

Temperature To convert from Celsius (°C) to Fahrenheit (°F), multiply by 1.8 and add 32. To convert from Fahrenheit (°F) to Celsius (°C), subtract 32 and divide by 1.8.

Appendix 2 Figure Legends

Igneous rocks and fossils

294 Appendix 2

Figure legends 295

Appendix 3 Geologic time-scale

Glossary

The age in years of a particular geologic event or feature, generally obtained with radiometric dating techniques. ABYSSAL (DEEP-SEA) PLAIN The ocean ﬂoor stretching from the continental margin to the mid-ocean ridge, usually very ﬂat with a slight slope. Sediments are generally thin and composed of the ﬁnest clastic material as well as true pelagic sediments. ACCOMMODATION SPACE The volume in a basin of sedimentation available for the accumulation of sediments. This space is created by sea-level rise, subsidence, or a combination of these two processes. ACOUSTIC IMPEDANCE The product of seismic velocity and density. Acoustic impedance often changes at the interface between two rock types, allowing a good seismic reﬂection to be returned. AIREY ISOSTASY A concept of the mechanism of isostasy in which the topographically higher mountains have the same density as other crustal blocks but have greater mass and deeper roots. ALLUVIAL FAN An assemblage of sediments marking the place where a stream moves from a steep gradient (e.g. in a mountain gorge) to a ﬂatter gradient and suddenly loses transporting power. Typical of arid and semi-arid climates but not conﬁned to them. ALLUVIAL PLAIN A level or gently sloping land surface, ABSOLUTE AGE

usually a ﬂoodplain or delta, produced by the extensive deposition of alluvium. AMMONITES A coiled, chambered fossil shell of a cephalopod mollusk of the extinct order Ammonoidea. ANDEAN BELT Magmatic arcs that develop on continental crust. They are presently only found along the west coast of Central and South America. ARAGONITE A mineral identical in composition to calcite, but differing from it in its crystalline form and some of its physical characters. It is not stable at the earth’s surface but despite this many organisms secrete aragonite skeletons and aragonite is a common early cement. ARCHITECTURAL ELEMENTS The minimum number of types of sedimentary units required to adequately deﬁne the process variation in a system. ASTHENOSPHERE The layer or shell within the earth beneath the lithosphere. It is a relatively weak zone in which isostatic adjustments take place and where magmas are generated. ASTROCHRONOLOGY A technique used to produce a geologic time-scale based on the correlation of sedimentary cycles, or other cyclic variations in paleoclimate with the earth’s orbital cycles. The evidence is best preserved in the climatic records of deep-sea sediments. BACKSTRIPPING A technique that uses the stratigraphic record to estimate quantitatively the depth that

298 Glossary basement would be in the absence of sediment and water loading. It is a quantitative way of working out when and where sediment accumulated and whether the sediment supply kept pace with subsidence and sea-level changes. BANDED IRONSTONE A rock that is made up of alternating light-colored, silica-rich layers and dark-colored layers of iron-rich minerals, which were deposited in marine basins on every continent about two billion years ago. BARRIER BAR A low, sandy island near shore and parallel to it, on a gently sloping offshore bottom. BASE LEVEL OF AGGRADATION The highest level at which a sedimentary deposit can be built. BASE LEVEL OF EROSION The lowest level at which a stream can erode the channel through which it ﬂows, generally equal to the prevailing global sea level. BASIN ANALYSIS A process that involves making an interpretation of the formation, evolution, architecture, and ﬁll of a sedimentary basin by examining geologic variables associated with the basin. Basin analysis provides a foundation for extrapolating known information into unknown regions in order to predict the nature of the basin where evidence is not available. This can assist the exploration and development of energy, mineral, and other resources that may occur within sedimentary basins. BEDFORM Any deviation from a ﬂat bed, generated by ﬂow on the bed of an alluvial channel. BIOCHEMICAL SEDIMENT A sediment resulting directly or indirectly from the chemical processes and activities of living organisms. BIOFACIES A subdivision of a stratigraphic unit distinguished from adjacent units on the basis of fossils alone. Biofacies can be used to reﬁne environments originally based on lithofacies. BIOSTRATIGRAPHY The branch of geology concerned with the separation and differentiation of rock units by study of the fossils they contain. BIOTURBATION The mixing and movement of sediment by organisms. BOREHOLE CHIPS Fragments of rock brought up from a known level during the drilling of boreholes. BOREHOLE CORE A cylindrical section of rock, usually 5–10 cm in diameter, produced from drilling. BOUGER GRAVITY ANOMALY A gravity anomaly calculated after corrections for latitude, altitude, and terrain. Bouger anomalies are highly negative in mountainous areas because the corrections do not take into account the low-density roots of the mountains.

A characteristic sequence that makes up a complete turbidite. From the bottom upward, this sequence includes (A) coarse graded bedding, (B) planar bedding, (C) current ripple laminations, (D) upper parallel laminations, and (E) clay. BRAID PLAIN A plain formed by a braided stream that has multiple channels caused by highly ﬂuctuating sediment-laden discharges. Often synonymous with a glacial outwash plain. CALCITE A common rock-forming mineral: CaCO3. Calcite can be white, colorless, or pale shades of gray, yellow, and blue. It readily effervesces (bubbles) in hydrochloric acid and is the principal component of limestone. CARBONATE SLOPE A carbonate depositional environment. Such slopes may be much steeper than clastic slopes because of the rapid growth of reefs as well as early cementation. CATASTROPHIC CHANGE The hypothesis that a series of immense, brief, worldwide upheavals changed the earth’s crust greatly and can account for the development of mountains, valleys, and other features of the earth. CHEMICAL SEDIMENT Sediment that is composed of previously dissolved minerals that have either precipitated from evaporated water or been extracted from water by living organisms and deposited when the organisms died or discarded their shells. CHEMICAL WEATHERING The process by which chemical reactions alter the chemical composition of rocks and minerals that are unstable at the earth’s surface and convert them into more stable substances. CHERT Granular cryptocrystalline silica, similar to ﬂint but usually light in color. Occurs as compact massive rock or as nodules. The origin of the silica is normally biological, from diatoms, radiolaria, or sponge spicules. CHRONOSTRATIGRAPHY The organization of rock strata into units based on their age or time of origin. CLASTIC SEDIMENT Sediment composed primarily of broken fragments derived from pre-existing rocks or sediments that have been transported some distance from their point of origin. CLIMATE The characteristic weather of a region, especially wind, temperature, and precipitation averaged over some signiﬁcant period of time. CLIMATIC EUSTASY Worldwide sea-level change related to a change in average global temperature. Sea-level change is accomplished by adding or removing water from the ocean, most commonly by melting or BOUMA SEQUENCE

Glossary 299 forming ice, or due to thermal expansion/contraction of the ocean water. COLLISION OROGENS Areas of extreme deformation where magmatic arcs collide with continents or where continents collide. COMPOSITE CYCLE A composite cycle is when cycles of differing wavelength are superimposed on one another. A good example would be the various Milankovitch climatic cycles superimposed on one another to give a very complex stratigraphic sequence. CONCORDIA PLOT The graph produced when the 238U/206Pb ratio is plotted against the 235U/ 207Pb ratio as both increase in value due to nuclear decay of uranium to lead over time in a closed system. CONTINENTAL RIFT Continental rifts form where continental crust is being pulled apart. A modern example of the initial stage of the process is the East African Rift system. CONTINENTAL SHELF Part of the continental margin. The shelf comprises the ocean ﬂoor from the coastal shore of continents to the continental slope, usually to a depth of about 100–250 m. The continental shelf usually has a very slight slope, roughly 0.1°. CONTINENTAL SLOPE The portion of ocean ﬂoor extending downward from the seaward edge of continental shelves. In some places, such as south of the Aleutian Islands, slopes descend directly to ocean deeps. In other places, such as off eastern North America, they grade into somewhat gentler continental rises, which in turn lead to deep ocean ﬂoors. Th continental slope coincides with the structural change from continental to oceanic crust. CONTOURITE Contourites usually consist of ﬁne sand or coarse silt deposited on the continental rise by contour-following bottom currents. CRATONIC BASIN A sedimentary basin on top of a craton. They tend to be bowl shaped with a long history of slow subsidence and have a ﬁll of thin but extensive sedimentary units. CYCLE An ordered succession of rock types that are normally recurrent and return to a starting point, that are repeated in the same order at more or less regular intervals and that end under conditions similar to those at the start of the cycle. The pattern does not have to be complete or fully developed in all cases. CYCLOTHEM A more complicated cycle typically associated with unstable shelf or interior basin conditions in which alternate marine transgressions and regressions occur. DEBRIS AVALANCHE (1)A turbulent granular ﬂow caused

by large landslides in which the particles are lubricated by the ambient ﬂuid and in which the larger clasts can impact during movement. (2) The very rapid and usually sudden sliding and ﬂowage of incoherent, unsorted mixtures of soil and weathered bedrock. DEBRIS FLOW A mixture of sediment and ﬂuid in which the larger clasts are supported by matrix strength and ﬂuid buoyancy as well as by the grain dispersal forces of debris avalanches. A distinctive characteristic is that the supporting ﬂuid does not mix with the ambient ﬂuid. DEEP-SEA FAN A terrigenous cone or fan-shaped deposit located seaward of large rivers or submarine canyons. It forms when turbid suspensions ﬂow down continental slopes and decelerate upon reaching the deepsea ﬂoor. Also known as submarine fans. DEFLATION The process by which wind erodes bedrock by picking up and transporting loose rock particles. DELTA A plain underlain by an assemblage of sediments accumulated where a stream ﬂows into a body of standing water, its velocity and transporting power suddenly reduced. Originally named after the Greek letter delta (D) because many are roughly triangular in plan, with the apex pointing upstream. DEPOSITIONAL SEQUENCE A relatively conformable succession of genetically related strata bounded by unconformities and their correlative conformities. DIAMICTITES A non-sorted or poorly sorted terrigenous sedimentary rock or sediment that contains a wide range of clast sizes (e.g. till or tillite). DIASTEM A relatively short interruption in sedimentation with little or no erosion before sedimentation resumes. Usually these breaks in the stratigraphic record are too short to be measured accurately except by paleontological evidence. DOLOMITE A mineral with the formula (Ca,Mg)(CO3)2. Also a sedimentary rock containing more than 50% of the mineral dolomite. DOLOSTONE A sedimentary rock composed primarily of the mineral dolomite. Dolostone is thought to form when magnesium ions replace some of the calcium ions in limestone, to which dolostone is similar in both appearance and chemical structure. DOWNLAP A stratigraphic relationship in which initially inclined strata terminate downdip against an initially inclined or horizontal surface. EPICONTINENTAL (EPEIRIC) SEA A sea on the continental shelf or within a continent. This type of sea covered vast areas of the continents in ancient times, for instance

300 Glossary the middle Ordovician sea of North America stretched almost unbroken across the entire continent. EPHEMERAL STREAM A stream or reach of a stream that only ﬂows brieﬂy in response to rainfall or melting snow or ice and whose channel is above the water table at all times. ERG A wind-deposited sand sea (an extensive area of sandy desert). EROSIONAL DIASTEM A diastem produced by a small amount of erosion rather that a short period of nondeposition. EUSTASY The worldwide sea-level regime and its ﬂuctuations caused by absolute changes in the quantity of sea water (e.g. by continental ice-cap ﬂuctuations). EUSTATIC (ABSOLUTE) SEA-LEVEL CHANGE A change in sea level produced entirely by an increase or decrease in the amount of water in oceans; hence of worldwide proportions. EVAPORITE A rock composed of minerals precipitated from solutions normally concentrated by the evaporation of sea water. Examples: rock salt, gypsum, anhydrite. EXOTIC TERRANE A large area of rock occurring in a lithological association foreign to that in which it originally formed. EXTRACLAST A fragment of material in a carbonate rock produced by erosion of an older rock outside the area of deposition. FACIES An assemblage of mineral, rock, or fossil features reﬂecting the environment in which the rock was formed. FACIES CHANGE A lateral variation in the lithological or paleontological characteristics of contemporaneous sedimentary deposits. It is caused by or reﬂects a change in the depositional environment. FACIES FOSSILS A fossil, usually a single genus or species, that is restricted to a deﬁned stratigraphic facies or is adapted to life in a restricted environment. It prefers certain ecological surroundings and may exist in them with little change for long periods of time. FISSION TRACK Marks left in the latticework of a uraniumbearing mineral crystal by subatomic particles released during the ﬁssion of a radioactive atom trapped inside the crystal. FLASER BEDDING Ripple cross-lamination in which mud streaks are preserved in the troughs but incompletely or not at all on the crests. FLOATING TEXTURE A clastic sedimentary texture in which the individual clasts are not touching one another but are supported by a ﬁner matrix.

A surface separating younger from older strata across which there is evidence of an abrupt increase in water depth. FOREBULGE A upwarping of the earth’s crust a short distance in front of the margins of a thick continental ice sheet. It is caused by movement of mantle material from beneath the ice sheet to lower-pressure areas adjacent to it. FORMATION The fundamental unit of lithostratigraphic classiﬁcation. Most formations are tabular in shape and are mappable at the earth’s surface at a scale of 1 : 50,000 or are traceable in the subsurface. FRAMEWORK TEXTURE A clastic sedimentary texture in which the individual clasts are touching and supported by one another to form a rigid structure. GAL A unit of gravitational acceleration equal to one centimeter per second per second (named after Galileo). GEOCHRONOLOGY The study of time in relation to the history of the earth, especially by absolute age determination and relative dating systems devised for this purpose. GEODESY That branch of applied mathematics which determines, by means of observations and measurements, the shape and areas of large portions of the earth’s surface, or the general shape and dimensions of the earth. Also that branch of surveying in which the curvature of the earth is taken into account, as in the surveys of States, or of long lines of coast. GEOID The theoretical shape of the earth produced such that the force of gravity acts perpendicular to the surface at all points. GEOPOTENTIAL The potential energy of a unit mass relative to sea level, numerically equal to the work required to lift the mass from sea level to its present elevation. GLACIAL STAGE A major subdivision of the glacial epoch, especially one of the cycles of growth and disappearance of the Pleistocene ice sheets. GRAPESTONE A cluster of sand-sized grains of various type (often pellets) cemented shortly after deposition. The grains are often micritized and cemented by aragonite. The cluster has a lumpy outer surface similar in appearance to a bunch of grapes. GRAPHIC CORRELATION A method whereby the ranges of fossils in two or more sections are correlated statistically. GRAVITY (1) The force of attraction exerted by one body in the universe on another. Gravity is directly proportional to the product of the masses of the two FLOODING SURFACE

Glossary 301 attracted bodies. (2) The force of attraction exerted by the earth on bodies on or near its surface, tending to pull them toward the earth’s center. GRAVITY ANOMALY Difference between observed and computed values of gravity. GROUND-PENETRATING RADAR A geophysical tool that uses radio frequency on the order of 10–1000 MHz to explore the shallow subsurface (less than 30 m under normal conditions). It is only of use in media without a signiﬁcant clay or salt-water content. GROUP The formal lithostratigraphic unit next in rank above formation. A group includes two or more contiguous or related formations with signiﬁcant features in common. HALF-LIFE The time needed for one-half of the nuclei in a sample of a radioactive element to decay. HARDGROUND Sediment at the sea bottom, usually only a few centimeters thick, which becomes lithiﬁed to form a hard surface that is encrusted, bored, discolored, and modiﬁed by solution. It implies a break in sedimentation and may be preserved stratigraphically as an unconformity. HIGHSTAND The interval of time during one or more cycles of relative sea-level change when sea level is above the shelf edge in a given area. HIGHSTAND SYSTEMS TRACT A systems tract bounded below by the downlap surface and above by the next sequence boundary. The early highstand commonly consists of an aggradational parasequence set; the late highstand is composed of one or more progradational parasequence sets. It forms during stagnant or slowly falling sea level. HUMMOCKY CROSS-STRATIFICATION A type of crossstratiﬁcation in which the lower bounding surfaces are erosional and commonly slope at angles of less than 10°. It is thought to be caused by the action of storm wave surged on the shoreface. HYPERCONCENTRATED FLOW A type of ﬂow intermediate between a non-turbulent debris ﬂow and a fully turbulent water ﬂow. A type of turbulent stream ﬂow in which the sediment load exceeds the theoretical capacity due to sediment bulking by bank slumping and the like. Hyperconcentrated ﬂows are ﬂuid enough and turbulent enough to let differential settling form graded beds. If enough sediment is incorporated, a debris ﬂow is generated. HYPSOMETRIC CURVE A cumulative frequency proﬁle representing the statistical distribution of the absolute or relative areas of the earth’s solid surface above or below a given elevation, usually sea level.

The scale may be planetary down to a single drainage basin. HYPSOMETRIC INTEGRAL An expression of the relative volume of a landmass at a given contour. HYPSOMETRY The science of determining, by any method, height measurements on the earth’s surface with reference to sea level. ICE-CONTACT DEPOSITS Stratiﬁed drift in contact with melting glacier ice. Eskers, kames, and pitted outwash are common examples. IMMATURE SEDIMENT Sediments that are little changed from the materials making up the parent rock. This implies a short time and distance of transport. IMPACT CRATER A crater formed on the surface of a planet or moon by impact of an unspeciﬁed object. INCISED STREAM A stream that has been cut or entrenched into a surface because of relative uplift of the surface. Anastomosing channels stabilized by vegetation are the norm. Sedimentation is normally by vertical accretion, with stacked shoestring channel sands passing laterally into alternating marsh and swamp clays and peat/coal. INCISED VALLEY FILLS The ﬁlling of valleys originally formed by ﬂuvial systems that extend their channels basinward and erode into underlying strata in response to a drop in relative sea level. INDEX FOSSIL The fossil of an abundant, distinctive, widely distributed organism known to have existed for a relatively short period of time, used to date the rock in which it is found. INTERGLACIAL STAGE A subdivision of the glacial epoch separating two glaciations, characterized by a relatively long period of warm or mild climate during which the temperature rose to at least that of the present day. INTERSTADIAL A subdivision within a glacial stage marking a temporary retreat of the ice. INTRACLAST A component of limestone representing a torn up and reworked fragment of an earlier sediment that has been eroded in the basin of deposition. The size may range from ﬁne sand to gravel. IRONSTONE A hard sedimentary rock rich in iron. Ironrich minerals include siderite, pyrite, goethite, hematite, glauconite, and chamosite. ISLAND ARC A chain of islands rising above oceanic crust behind subduction zones (e.g. the Aleutians). This often happens at the edge of oceans where the lithosphere tends to be oldest and thickest. ISOPACH An isoline that connects points of equal thickness of a geologic stratum formation or group of formations.

302 Glossary Ideal condition of balance that would be attained by earth materials of differing densities if gravity were the only force governing heights relative to each other. LENS A geologic deposit bounded by converging surfaces. It is thick in the middle, and thins and ﬁnally disappears toward both edges. LITHOFACIES A lateral, mappable subdivision of a designated stratigraphic unit, distinguished from adjacent subdivisions on the basis of lithology. Lithology includes all mineralogical and petrographic characteristics as well as paleontologic characteristics that affect the composition, appearance, or texture of the rock. LITHOSPHERE The rigid outer layer of the earth, which includes the crust and upper part of mantle. Relatively strong layer in contrast to the underlying asthenosphere. LITHOSTRATIGRAPHY The element of stratigraphy that deals with the lithology of strata and their organization into units based on lithologic character. LOESS Unconsolidated, unstratiﬁed aggregation of siltsized, angular mineral fragments, usually buff in color. Generally believed to be wind-deposited and characteristically able to stand on very steep to vertical slopes. LOWSTAND The interval of time during one or more cycles of relative sea-level change when sea level is below the shelf edge. LOWSTAND SYSTEMS TRACT The lowstand systems tract is bounded below by the type 1 sequence boundary and above by the ﬁrst major ﬂooding surface, called the transgressive surface. It may consist of a lowstand fan and turbidites. MAGNETOSTRATIGRAPHY All parts of stratigraphy based on paleomagnetic signatures. These are normally the characteristic patterns formed when the earth’s magnetic ﬁeld reverses polarity (each reversal is thought to take about 10,000 years). MATURE SEDIMENT A clastic sediment that has differentiated or evolved from its parent rock by processes acting over a long time and with high intensity. These sediments are characterized by stable minerals such as quartz, deﬁciency of mobile oxides, absence of weatherable material, and well-sorted but angular to subangular grains. MATURITY The extent to which a clastic sediment texturally and compositionally approaches the ultimate end product to which it is driven by processes acting on it. ISOSTASY

A sequence of condensed facies that separates the transgressive systems tract from the overlying regressive highstand systems tract (see Flooding Surface). MECHANICAL WEATHERING The process by which a rock or mineral is broken down into smaller fragments without altering its chemical makeup. MEMBER A lithostratigraphic unit of subordinate rank comprising a laterally persistent rock unit within a formation. MICRITIZATION A decrease in the size of sedimentary carbonate particles, possibly due to boring algae. MID-OCEAN RIFT A valley that has developed along the center of a mid-ocean ridge system where oceanic crust is being pulled apart. Chemical reactions with hot circulating brines modify the basalt lava and the sediments along the rift. Pelagic sediments such as radiolarian chert and foraminiferal ooze may blanket the system. MILANKOVITCH CYCLE An explanation for climate change that results from the ﬂuctuations in the seasonal and geographic distribution of insolation. This is determined by the earth’s orbital elements, namely, eccentricity, tilt of the rotational axis, and longitude of perihelion. It is supported by radiometrically dated reconstructions of ocean temperature and glacial sequences. NON-DEPOSITIONAL DIASTEM A relatively short interruption in sedimentation with little or no erosion before sedimentation resumes. NUMERICAL TIME The ﬁxing of a geologic structure or event in time, by counting tree rings or radiometric dating, for example. OCEAN FLOOR Those parts of ocean basins that are underlain by oceanic crust. Excluded are the continental shelf and epieric seas. Ocean ﬂoor generally consists of an ophiolite basement overlain by a variety of volcanic and sediment types. ONCOIDS An algal biscuit that resembles an ancient oncolite. It is smaller than a stromatolite and generally is smaller than 10 cm in diameter. ONLAP An overlap generally characterized by the pinching out toward the margins of the basin of deposition of the sedimentary units within a conformable sequence of rocks. It shows successive layers extending up an inclined surface and usually indicates a transgression. OOID (OR OOLITH) Spheroidal grains of sand size, usually composed of calcium carbonate, CaCO3, and thought to have originated by inorganic MAXIMUM FLOODING SURFACE

Glossary 303 precipitation. Some limestones are largely made up of oolites. OUTCROP The part of a rock formation that appears at the surface of the ground. PALEOCURRENT A current, usually in water, of the geologic past. Evidence of the current direction, depth, and ﬂow regime is recorded in sedimentary rocks. PALEOMAGNETIC POLE The exact location to which a compass needle points. This location changes over time and ancient positions of the earth’s poles can be determined based on paleomagnetic evidence. PALEOTHERMOMETRY Measurements or estimates of paleotemperature, especially the temperature determinations of shallow seas based on oxygen isotope analysis. PARASEQUENCE A relatively conformable succession of genetically related beds or bedsets bounded by ﬂooding surfaces or their correlative surfaces. In special positions within the sequences, parasequences may be bounded either above or below by sequence boundaries. PARASEQUENCE SET A succession of genetically related parasequences forming a distinctive stacking pattern bounded by major ﬂooding surfaces and their correlative surfaces. PASSIVE MARGIN A continental margin that does not coincide with a plate margin and has a minimum of tectonic activity as a result. It consists of slowly subsiding and thinning continental crust at the edge of an opening ocean. PEAT Partially reduced plant or wood material, containing approximately 60% carbon and 30% oxygen. An intermediate material in the formation of coal. PELOIDS An allochem composed of micrite, irrespective of size or origin. The term includes pellets and intraclasts, and is useful where the exact origin is unknown. PHOSPHATE A general term used to describe sedimentary rock containing calcium phosphate along with halides and hydroxides. Phosphates form by accretion as nodules in the deep sea and as replacements of other minerals in shelf and slope environments. PHREATIC Relating to groundwater or relating to an explosion caused by steam derived from groundwater. PHYSICAL CORRELATION This involves tracing some physical aspect or relationship of the strata such as an unconformity, ash layer, or a unique rock or fossil type from one area to another. PHYSICAL WEATHERING The breaking of rocks and minerals at the earth’s surface into smaller pieces by mechani-

cal action. This action may include frost wedging, sandblasting, abrasion, plucking, and other processes. PISOID A small, rounded accretionary body in a sediment, resembling a pea in size and shape. It is often formed of calcium carbonate. PLAYA Flat-ﬂoored center of an undrained desert basin. Playas are often sites for the deposition of evaporite minerals. PRATT ISOSTASY A concept of the mechanism of isostasy that postulates an equilibrium of crustal blocks of varying density. Thus the topographically higher mountains would have a lower density than topographically lower units, and the depth of crustal material would be the same everywhere. PRINCIPLE OF CROSS-CUTTING RELATIONSHIPS The scientiﬁc law stating that an intrusion (both igneous and sedimentary) or fault is always younger than the rock that surrounds it. PRINCIPLE OF FAUNAL SUCCESSION The scientiﬁc law stating that speciﬁc groups of animals have followed, or succeeded, one another in a deﬁnite sequence through earth history. PRINCIPLE OF ORIGINAL HORIZONTALITY The scientiﬁc law stating that sediments settling out from bodies of water are deposited horizontally or nearly horizontally in layers that lie parallel or nearly parallel to the earth’s surface. PRINCIPLE OF SUPERPOSITION The scientiﬁc law stating that in any unaltered sequence of rock strata, each stratum is younger than the one beneath it and older than the one above it, so that the youngest stratum will be at the top of the sequence and the oldest at the bottom. PRINCIPLE OF UNIFORMITARIANISM The scientiﬁc law stating that the geologic processes taking place in the present operated similarly in the past and can therefore be used to explain past geologic events. PROVENANCE The area from which the constituents of a sedimentary rock or facies are derived. If this can be determined then inferences can be made about the erosion and transport of the sediment. RADIOMETRIC DATING Calculating an age in years for geologic materials by measuring the presence of a shortlife radioactive element or by measuring the ratio between a long-life radioactive element and its decay product. RAMP A type of carbonate deposition environment in an area of the continental shelf that deepens toward the continental slope and is swept by waves because there is no marginal barrier.

304 Glossary A submerged ridge of rock or coral near the surface of the water. To a geologist the term means a massive three-dimensional carbonate structure surrounded by bedded carbonate rocks. REFERENCE SPHEROID A theoretical ﬁgure whose dimensions closely resemble those of the geoid and whose exact dimensions are determined by various properties of the section of the earth’s surface concerned. RELATIVE SEA-LEVEL CHANGE Sea-level change relative to a predetermined datum on the sea ﬂoor. Depositing sediment on the datum does not change the distance between the datum and the sea surface, hence deposition alone cannot cause a relative change in sea level. RELATIVE TIME Dating of events by placing them in chronological order of occurrence rather than in years. Compare with absolute time. RESISTIVITY LOG A well log produced when a current is passed through borehole walls by two contact electrodes and potential differences are measured between two or more electrodes. Deﬂections on the resistivity logs can be calibrated to known lithologies. RIMMED SHELF A type of carbonate deposition environment in an area of the continental shelf that has coral reefs, sand shoals, or islands along the seaward margin. RIVER-DOMINATED COASTLINE A deltaic coast characterized by lobate seaward bulges crossed by river distributaries and bordered by lowlands. SABKHA A supratidal environment of deposition on coastal plains just above the high-tide mark under arid conditions. They are characterized by salt, tidal ﬂood, and eolian deposits. SEDIMENTARY BASIN A low area in the earth’s crust, often of tectonic origin, in which sediments accumulate. Such features were basins at the time of deposition but are not necessarily so today. SEISMIC STRATIGRAPHY The study of stratigraphy and depositional facies as interpreted from seismic data. SEQUENCE A relatively conformable sequence of genetically related strata bounded above and below by unconformities or their correlative conformities. Parasequences and parasequence sets are the building blocks of the sequence. SEQUENCE STRATIGRAPHY Sequence stratigraphy uses the geometric relationships of strata on seismic sections to subdivide and correlate units. SERIR A stony desert plain strewn with rounded pebbles from which the sand has been blown away. REEF

SHELF-MARGIN SYSTEMS TRACT

This may form as a prograding wedge below the shelf edge. SOIL CATENA A sequence of soils of about the same age, derived from similar parent material, and occurring under similar climatic conditions, but having different characteristics due to variations in relief and drainage. SOIL HORIZON Layer of soil approximately parallel to the land surface with observable characteristics produced through operation of soil-building processes. SPONTANEOUS POTENTIAL An electrical exploration method in which one determines the electrical potentials caused by electrochemical reactions associated with clay or metallic mineral deposits. Also known as self-potential. STADIAL A substage of a glacial stage marked by a readvance of the ice. STATIC CORRECTION A technique used to normalize seismic information to a datum. It is used to overcome elevation differences between the source and the receiver, variations of water depth, etc. STRATA A sheet-like mass of sedimentary rock or earth of one kind lying between beds of other kinds, or a layer in which archaeological material (as artifacts, skeletons, and dwelling remains) is found on excavation. STRATIFICATION Structure produced by deposition of sediments in layers or beds. SUBDUCTION ZONE Elongate region along which thickened oceanic lithosphere starts gravitationally sinking beneath other oceanic or continental lithospheric blocks to form deep ocean trenches. SYSTEMS TRACTS Linkage of contemporaneous depositional systems composed of parasequences and one or more parasequence sets forming the subdivisions of a sequence. Lowstand, transgressive, shelf margin, and highstand systems tracts are recognized. Systems tracts are interpreted based on parasequence stacking patterns, positions within the sequence, and types of bounding surfaces. TAPHONOMY The study of what happens to a fossil, from the time of its initial creation (e.g. the death of an organism or the imprint left by the movement of an organism) to the time that the fossil is discovered by a paleontologist. For example, shells or bones can be moved by running water, and later be compressed by overlying sediment. Taphonomy is often broken into two parts, biostratinomy and the study of diagenesis. TECTONIC EUSTASY Global sea-level change caused by changes in the capacity of the oceans. The capacity can be altered by increasing or decreasing the

Glossary 305 area and/or depth of the oceans through tectonic activity. TILL Unstratiﬁed, unsorted, glacial drift deposited directly by glacier ice. TONGUE A geologic deposit bounded by converging surfaces. It is thick at one end and dies out laterally in one direction. TOPLAP Shows inclined layers prograding outward and asymptotically into an overlying surface. It usually indicates a subaqueous non-depositional surface of a sedimentary deposit. TRANSFORM BELT A narrow, complex, fault-controlled basin or series of basins with dominantly strike-slip faulting. This happens at plate boundaries that have mostly lateral movements. TRANSGRESSIVE SYSTEMS TRACT A systems tract bounded below by the transgressive surface and above by the downlap surface or maximum ﬂooding surface. Parasequences within the transgressive systems tract backstep in a retrogradational parasequence set. The systems tract progressively deepens upward as successively younger parasequences step farther landward. The downlap surface, coincident with the upper boundary of the youngest parasequence in the transgressive systems tract, is the surface onto which the clinoform toes of the overlying highstand systems tract may merge and become very thin. TREE RINGS Annual growth rings seen in the crosssection of a tree trunk. The pattern of rings reﬂects the climate variations during growth, so that by comparing and overlapping the pattern of rings of dead and living trees a dendrochronology can be established. TURBIDITY CURRENT A current in which a limited volume of turbid or muddy water moves relative to the surrounding water because of the current’s greater density. TYPE 1 SEQUENCE A sequence composed of lowstand, transgressive, and highstand systems tracts bounded beneath by a type 1 unconformity and the correlative conformity. Depending upon where the sequence is observed in the basin, not all of the systems tracts may be present due to truncation or non-deposition. This type of sequence is interpreted as forming when the rate of eustatic fall exceeds the rate of subsidence at the depositional–shoreline break, producing a relative fall in sea level at that position. TYPE 1 SEQUENCE BOUNDARY A boundary showing subaerial erosion with channels incised into the shelf and deposition of sediment at the base of the slope. It is

caused by the base level dropping below the shelf edge. TYPE 2 SEQUENCE A sequence composed of shelf-margin, transgressive, and highstand systems tracts bounded beneath by a type-2 unconformity and the correlative conformity. A type-2 sequence boundary is interpreted as forming when the rate of eustatic fall is slightly less than or equal to the rate of basin subsidence at the existing depositional–shoreline break at the time of the eustatic fall. This means that there is no relative fall in sea level at the depositional–shoreline break for the type-2 sequence. Exxon no longer recognizes a type-2 sequence. Presently, a type-2 sequence boundary is considered synonymous with a parasequence set boundary, and a type-2 sequence is considered synonymous with linked progradational and retrogradational parasequence sets. TYPE 2 SEQUENCE BOUNDARY A boundary lacking extensive subaerial erosion and valley incision since the base level does not drop below the shelf edge. TYPE SECTION The originally described sequence of strata that comprise a formation. Normally this is in an area where the formation is particularly well exposed and where its character is seen. It serves as an objective standard to which other parts of the formation can be compared. UNCONFORMITY A buried erosion surface separating two rock masses, the older rocks having been exposed to erosion for a long interval of time before deposition of younger rocks. If the older rocks were deformed and not horizontal at the time of subsequent deposition, the surface of separation is an angular unconformity. If the older rocks remained essentially horizontal during erosion, the surface separating them from younger rocks is called a disconformity. An unconformity that develops between massive igneous or metamorphic rocks exposed to erosion and then covered by sedimentary rocks is called nonconformity. VARVED CLAY A ﬁne-grained deposit containing couplets of thin sedimentary beds, one coarser and thicker, one ﬁner and thinner. This couplet is interpreted as representing a cycle of one year in lakes fringing a glacier, the thicker layer representing summer and the thin layer winter. WALTHER’S LAW The law stating that in a continuous vertical sequence, the succession of facies vertically reﬂects the original lateral arrangement of facies in the area. To put it another way, the sequence of environments that can be seen laterally on the earth’s surface can also be seen vertically in a pit or well dug

306 Glossary through the sequence. This is especially well seen in prograding shorelines, deltas, point-bar deposits, and the like. WEATHERING The response of materials once in equilibrium within the earth’s crust to new conditions at or near contact with water, air, or living matter. WIRELINE LOG A graphic record of measurements made down a well or borehole. Several measurements are

usually made at once and the resulting plots are displayed side by side on a common vertical scale. ZONE A minor interval in any category of stratigraphic classiﬁcation. There are many kinds of zones: biozones, chronozones, lithozones, mineral zones, etc. The term should always be used with a modiﬁer. ZONE FOSSIL A fossil characteristic of a certain time regardless of depositional environment.

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Index

Note: Page numbers in italics refer to ﬁgures and plates and those in bold to tables. absolute sea-level changes see eustatic sealevel changes Abukuma (Japan), 231 abundance zones, 146 abyssal plains, 91–3, 127 accommodation space, 192–6 acme zones, 146 acoustic impedance, 119, 121 Acropora spp. (corals), 51 Aden, Gulf of, 222, 225 aerial photography, 9 Afar triangle, 225 Africa, 15, 243, 255 East, 222, 268 mammals, 278 West, Plate 8.3 AFTA (apatite ﬁssion track ages), 183 age equivalence, 116 concept of, 142 imprecise, 142 precise, 142 Ager, D.V., 129, 199, 205, 272 ages in years see numerical dating methods agglomerates, 40 aggradation, 106–9 aggregates, 50, 51 formation, 54 aggregation carbonates, 24 mechanisms, 20 Aigner, T., 9 air, threshold stresses, 38 air ﬂows, 22, 25, 30–2 air guns, 119 Airey isostasy, 214 Alaska (USA), 36, 91, 263

Albee, A.L., 287 Alberta (Canada), 1 Albritton, C.C. Jr., 140 algae, 47, 53, 58, 63 evolution, 267, 283 green, 50, 51, 54 red, 50 see also cyanobacteria algal mats, 94, 95 Alger, R.P., 124 alkaline earth carbonates, 60 Allen, J.R.L., 20, 22, 24, 67, 171, 221 Allen, P.A., 171, 215, 221 allocyclic processes, 189, 190, 192 allostratigraphic units, 201 alluvial fans, 69, 70, 75, 76, 235 arid, 78, 85 breccias, 175 deposits, 78 lateral changes, 127 alluvial plains, 73–4 Alps, 187 alteration, 36 mechanisms, 36 aluminum oxide, 20 Alvarez, Luis W., 25, 67, 129, 170 amalgamation, 235 Amazon Basin, 15 Amazon, River, 91, 210 ammonites, 5, 110, 170 evolution, 147 Lower Jurassic, 254 zones, 144, 145, 147, 148–9, 151 amphiboles, 166 amplitude maps, Plate 8.2 Anabaena spp. (cyanobacteria), 273

Andean belts, 228, 231–2 characteristics, 232 Anderson, E.J., 196 Andes, 91, 222 Angara terrane, 274 angiosperms, 262 angular unconformities, 106 anhydrite, 60, 62 animals fossil, 263 multicellular, 267 Annan, A.P., 281 Annan (Scotland), 37 anorthosites, 285 Anstey, N.A., 118 Antarctica, 15, 59, 79, 87–8 corals, 264 evaporites, 260 glaciation, 265 ice, 154, 156 anticlines, 122, 130 apatite, 168 apatite ﬁssion track ages (AFTA), 183 Apennines (Italy), 219, Plate 8.4 Appalachia (North America), 184 Appalachian–Caledonian structural belts, 275 Appalachians (USA), 7, 132, 184 apparent truncation, 172 Aptian, 197 Arabian Sea, 268 aragonite, 47, 50, 54, 55–6, 57–8, 153 dissolution, 58 precipitation, 59, 60 secretion, 51, 272

322 Index Archean, 59, 208, 216, 265, 283 atmospheric changes, 271 basin formation, 220 chronostratigraphy, 284 plate tectonics, 284 strata, 271 stratigraphy, 276, 282, 283, 284, 291 archeological stratigraphy, 276, 281–3 calibration, 281 issues, 282–3 archeology, dating errors, 168 Archimedes’ principle, 214 architectural elements, 70 deﬁnition, 67 Arctic, 262 Arduino, Giovanni (1714–95), 187 arenites, 42, 45 argon, radioactive decay rates, 165 argon–argon dating method, 166, 170, 279 argon isotopes, inherited, 166 Aristotle (384–322 BC), 140, 202 Arkell, William Joscelyn (1904–58), 145, 198 arkosic arenites, 42, 45 arkosic sandstones, 43, 44 Asia, 222, 268, 275 evaporites, 260 see also Central Asia; Southeast Asia assemblage zones, 146 Asteroids, 286 asthenosphere, 215 astrochronology, Quaternary, 278–9 astronomical events, 160 Aswan (Egypt), 111, 112 Atlantic Ocean, 147, 184, 185, 231 deep-sea plains, 93 Mid-Atlantic Ridge, 159, 213 opening, 275 trilobites, 274 see also North Atlantic atmosphere, 255–7, 283, 284 anoxic, 276 chemical changes, 152, 153–4 circulation systems, 268 evolution, 272 and greenhouse gases, 267–8 Mars, 272, 287 organism-induced changes, 271 stable isotopes, 263 Aubrey, M.-P., 201 Auk Field (North Sea), 120 aulacogens, 222, 225 Aulacostephanus spp. (ammonites), 147 Australia, 19, 78, 80, 231 carbonates, 50 cratons, 282 stromatolites, 273 autocyclic processes, 189, 190, 192

avalanches, 27, 114 deposits, 26 see also debris avalanches Avermann, M., 26 back-arc basins, 230, 231 backstripping, 175–80, Plate 8.5 backstripping equation, 178–9 bacteria, 20, 47, 58 see also cyanobacteria Bactrian kings, 283 Bafﬁn Island, 245 Bahamas, 54 Baigang Formation (China), 175 Baise Basin (China), 174–5 Baker, 184 ball-and-pillow structures, 33, 39 Bally, A.W., 120 Baltic Sea, 97, 99 banded ironstones, 100, 271, 283, 284 Bangka shear (China), 240 Baralacha La (India), 29 Barbados, 243 barchans, 39 Barnes, J.W., 134 Barrell, Joseph, 109, 190 barrier bars, 94 marine, 1, 2, 3 barrier coasts, 83–4, 87 barrier complexes, 84, 87, 88 barrier–lagoon complexes, 133, 136 barriers, 74 Barrovian metamorphism, 232, 234 bars, 70, 81, 95 oolitic, 94 point, 67, 73 sand, 83 see also barrier bars basalt dikes, 142 basalt lavas, 170, 223 basaltic shield volcanoes, 223 basalts, 21, 141, 170, 267 komatiitic, 284 paleomagnetic reversals, 279 see also ﬂood basalts Basilevsky, A.T., 291 basin analysis, 9, 171–85, 192 basin development, 210 and plate tectonics, 220–35 basin formation, 171, 173, 224 by vertical movements, 218–20 basin-ﬁll architecture, 171–4 basins, 196 back-arc, 230, 231 backstripping, 175–80 classiﬁcation, 219 evaporitic, 96 ﬁlling, 192 foreland, 232, 233, 234 impact, 220, 221 intermontane, 217

marginal, 232 remnant, 232 retro-arc, 231 rift, 175, 240 sediment thickness, 224 strike-slip, 232–3 types of, 230–1 see also cratonic basins; sedimentary basins Bassett, M.G., 105 Bathurst, R.G.C., 47 Bayer, U., 65 beach sands, 43, 131 beaches, 74 Beatty, J.K., 291 Becquerel, Antoine Henri (1852–1908), 8, 141 bedforms, 33, 39, 70 current-wave, 29 migration, 24, 28–9 beds, 201 boulder, 25 graded, 28, 190 lava, 200 organism-induced effects, 272 red, 260 shell, 83 belemnites, evolution, 147 Bellerophon spp., 133 benthic communities, 52 bentonite correlations, 126, 129 Bermuda, 95 Berry, W.B.N., 189 berthierine, 65 bicarbonates, 20 bimammatum zone, 131 biochemical sediments, 22, 46–65, 223 deposition, 23–4 formation, 23 occurrence, 67 bioclastic sediments, 94 changes, 271–2 biofacies, 131, 132, 133–4, 172, 202 archeological, 282 changes, 222 biogeographic changes, 272–5 biogeography, of organisms, 149 bioherm, use of term, 96 biology, 9, 271–5 biomes, 259 biosparite grainstones, 55 biostratigraphic classiﬁcation, 103, 104, 149 graphic correlation method, 149–52 biostratigraphy, 143–52, 158, 187, 201–2, 203, 276 approaches, 4–5 biological changes, 271 calibration, 157 concept of, 189 deﬁnition, 143

Index 323 Neoproterozoic, 283 Quaternary, 277–8 biotas biogeographic changes, 272–5 climatic effects, 262–3 multicellular, 276 biotite, 165, 166 age spectra, 167 bioturbation, 33–6, 42 mechanisms, 272 birdfoot deltas, 83, 84, 133 birds, fossil, 148 Bischke, R.E., 134 bivalves, 75 bivariate plots, 149 Blackwell, B.A., 281 Blatt, H., 22, 36, 102 Blue Lias, 191 blue-green algae see cyanobacteria Boltwood, Bertram Borden (1870–1927), 142, 162 borehole chips, 123 borehole cores, 123 boreholes, 1, 120, 122, 126, 171, 174 subsidence curves, 179, 181 thermal histories, 183 wireline logs, 123–5 Bork, K.B., 187 Bouger correction, 211–12 Bouger gravity anomalies, 212, 213 boulder beds, 25 Bouma sequences, 28, 31 Bowen’s reaction series, 13, 14 Boyd, R., 197 brachiopods, 51, 53, 133, 144 Brahma, 140 braid plains, 68, 70–2 characteristics, 76 braided streams, 68–9, 70, 272 Brand, U., 47, 48 breccias, 25, 26, 27, 285–6 formation, 78, 96 in impact craters, 220 Permian, 175, 176 types of, 40 Brecon Beacons (England), 105 Brett, C.E., 7 Bridge, J.S., 99 Bridport Sandstone (England), 114 Bright Angel Shale (USA), 132 brine shrimps, 94 Brinkmann, R., 110 bristle-cone pines, 160 British Columbia (Canada), 30, 31 Brockmeyer, P., 26 Bromley, R.G., 33, 272 Brookﬁeld, M.E., 7, 18, 196, 240 bryozoa, 50 Buffon, Georges-Louis Leclerc, Comte de (1707–88), 141 Bugge, T., 27

Burchﬁeld, J.D., 141 burrowing depth, 272 Burton, E.A., 50 Busby, C.J., 221 Bustin, R.M., 63 Butler, R.F., 155 Byers, C.W., 189 Calabria (Italy), 143 calcareous oozes, 50 calcareous rocks, 49 calcite, 47, 54, 100, 153 chemical decomposition, 14 formation, 55–6 forms, 47 precipitation, 47, 50 secretion, 272 see also high-magnesium calcite; lowmagnesium calcite calcium carbonate, 47, 50, 58 precipitation, 57, 59 calcium cations, 47, 49, 165 calcium ions, 49 calcium isotopes, 165 calcium–magnesium carbonate, 47 calcium phosphate halides, 65 calcium phosphate hydroxides, 65 calcium sulfate, 59, 62 calcium-sodium plagioclase, 36 calcretes, 20, 54, 94 formation, 15, 18 Caldeira, K., 285 calderas, 25, 26, 27 Caledonian, 235 calibration, 157, 279 archeological stratigraphy, 281 coins, 280, 283 dendrochronology, 168, 281 numerical dating methods, 170, 283 radiocarbon dating, 160, 168, 170, 281 caliche, 73 California (USA), 235–6, 263 California, Gulf of, 226, 234 Caloris Basin (Mercury), 287, 289 Cambrian, 7, 103, 105, 118, 132–3, 283 clays, 106 divisions, 104 faunas, 275 multicellular animals, 267 polar wanderings, 158 sediments, 172 trilobites, 274 Cambrian–Silurian boundary, 4–5 Canada, 1–3, 4, 215, 283 debris ﬂows, 30, 31 deltas, 246 peats, 62, 262 Quaternary, 277 sandstones, 177 sediments, 6

see also Hudson Bay (Canada); Ontario (Canada) Canadian system, 7 canals, 143, 144 canyons, 111, 112, 114, 116, 132–3 facies changes, 128 strata, 118 El Capitan reef (USA), 98, 127–8 Caradocian Series (England), 137 carbon, 25, 263 organic, 166, 263, 267 see also hydrocarbons carbon–carbon ratios, 166–8, 278 carbon dioxide, 46–7, 49, 50 exchange, 271 and greenhouse gases, 267–8 ions, 14 reduction, 269 removal, 59, 268 variations, 267 carbon dioxide isotopes, 166 carbon isotopes, 48, 49, 142, 202, 278 half-lives, 163 carbon-14 dating method, 142, 163, 166–8, 278 see also radiocarbon dating carbon-14 isotopes, radioactive decay rates, 166–8 carbonate banks, 95 carbonate–clay couplets, 190 carbonate muds, 53, 54, 131 carbonate reservoirs, 120 carbonate sediments, 81, 93–4 carbonate shelves, 94–6 carbonate shorelines, 94, 95 carbonate slopes, 96 carbonate systems, 93–7 and clastic systems compared, 93–4 occurrence, 93 carbonates, 18, 46–59, 260 alkaline earth, 60 cementation, 57–8 chemistry, 49–50 classiﬁcation, 46, 55 clastic, 46 concretion, 42 constituents, 53–4, 55 deep-sea, 96–7 deposition, 54–5 diagenesis, 49, 55–9 diagenetic segregation, 114 dissolution, 58 ferroan, 47 formation, 223 grain compaction, 58 in island arcs, 231 isotopes, 47–8, 263 isotopic composition, 154–5 lake sediments, 78 lithiﬁcation, 58 marine sediments, 50–3

324 Index carbonates (cont.) mineralogy, 47, 48 neomorphism, 58 Ordovician, 122 particle aggregation, 24 in passive margins, 228 pelagic, 153 in porewaters, 272 precipitation, 36, 46, 114–15, 154: freshwater, 49; salt-water, 50 production rates, 93 recycling, 99 replacement, 58–9 supply, 190 transportation, 54–5 weathering, 268 see also calcium carbonate carbonation, mechanisms, 14–15 Carboniferous, 1, 48, 63, 185, 272 coals, 108, 268 conglomerates, 177 ice ages, 267 swamps, 262 weathering, 271 Caribbean, 27, 99 Carmel Formation (USA), 114 Carney, J.L., 150 Carnian, 179 Carr, M.H., 287 Carter, R.M., 190 Cas, R.A.F., 26, 67 Caspian Sea, 81, 99 catastrophic changes, 111 catastrophism, 129–30, 187 catenas, soil, 20, 21 Catskill Delta (USA), 132–3 Catskill Sea, 97–9 cave pearls, 49 caves, 280, 282 Cecil, C.B., 262 cementation, 14–15, 36, 96, 272 carbonates, 57–8 conditions, 58 mechanisms, 36 nodular, 36 cements, 42 precipitation, 39 Cenozoic, 170, 192 basins, 222 clays, 170 divisions, 187 ﬂoras, 262 fossil plants, 262 mollusks, 263 rift basins, 240 sea-level changes, 243, 252 sediments, 171 temperature, 263 Central America, 231 Central Asia, 111, 283 basins, 81

deserts, 15 glaciers, 81 Central India, 18 Central Paciﬁc, 157 cephalopods, 5 cereals, 282 chalks, 93, 120, 129, 272 Chamley, H., 67 chamosite, 64, 65 Chandeleur Island (Gulf of Mexico), 88 Chateau du Joux (France), 130 Chave, K.E., 50 chemical changes, 152–4 chemical sediments, 22, 46–65 deposition, 24 formation, 22, 78, 223 occurrence, 67, 81 chemical weathering, 46 mechanisms, 13–14 rates, 15 chemostratigraphy, 203 disadvantages, 284 chemozones, 203 cherts, 18, 46, 59, 284 formation, 223 radiolarian, 93, 223 Chicxulub (Mexico), 170, 220, 221 China, 82, 174–5, 202, 232 China Basin, 240 chlorapatites, 65 chlorides, 59 chlorite, 45 Christ, Jesus, 140 Christianity, 187 Christiansen, E.H., 285 Christie-Blick, N., 192 chronostratigraphy, 103, 199–200, 201–3, 276, 282 Archean, 284 Quaternary, 277, 278 rock units, 201 chute-and-pool sands, 73 clams, 53 clarain, 63 classiﬁcation, 4, 8, 12 basins, 219 carbonates, 46, 55 chronostratigraphic, 103 cycles, 190 development, 186–9 Folk’s, 55 geochronologic, 103 lithostratigraphic, 103, 104 sandstones, 36, 44 sedimentary basins, 221 sediments, 36–9 spatial, 103 stratigraphic, 103–5 temporal, 103 Werner’s, 187

see also biostratigraphic classiﬁcation; Dunham’s classiﬁcation clastic sediments, 22, 24–46, 53, 55 accumulation, 177–8 characteristics, 36–9 deposition, 22–3, 24–5 formation, 22–3 fossils, 144 grain size, 42 maturity concept, 24 occurrence, 67, 83, 91 post-depositional modiﬁcations, 32–6 terminology, 39–46 clastic shorelines, 83 clastic systems, 69, 82–93 and carbonate systems compared, 93–4 clastics, 65, 184 marine, 234 clasts, 27, 54, 175, 176 see also extraclasts; intraclasts clay minerals, 14, 18, 21, 24, 45, 46 clays, 23, 29, 69, 84, 87, 127 analyses, 170 Cambrian, 106 compaction, 33, 114 delta, 114 deposition, 36 deposits, 73, 262 formation, 74, 78 lake, 67 marine, 246 offshore, 131 porosity, 36 prodelta, 83 Quaternary, 106 supply, 190 water loss, 41 see also varved clays climate, 9, 226, 255–69 and basin formation, 171 characteristics, 255–9 controls, 265–9 deﬁnition, 255 effects: on biotas, 262–3; on sediments, 259–62; on stable isotopes, 263–5 and environmental change, 67 and eustatic sea-level changes, 268–9 ﬂuctuations, 255 mechanisms, 255 paleoclimate indicators, 62 and plate tectonics, 268–9 present distribution, 255–9 and sea level, 270 and volcanism, 270 climate cycles, 190 climate regimes, and weathering, 15, 45 climatic changes, 75, 111, 278 climatic cycles, 190–2 climatic eustasy, 242, 243, 245–6 clocks, pendulum, 209

Index 325 coalﬁelds, 187 coaliﬁcation, 181 coals, 62–3, 63, 192, 193, 262, 268 Carboniferous, 108, 268 constituents, 63 formation, 272 lithotypes, 63 Pennsylvanian, 73–4 quartzites in, 271 coastal systems, classiﬁcation, 83 coastlines open, 94 river-dominated, 83 tide-dominated, 83 wave-dominated, 83–4 see also shorelines coated grains, 54 coins, 282–3 calibration, 280, 283 collision orogens, 232–4, 235 Colorado (USA), 193 Colorado Delta (USA), 85, 226 Colorado River (USA), 118 compaction, 39, 177, 178, 179 mechanisms, 36, 41 complexes, 201 composite standard sections, 150 concordia diagrams, 163–4 concretion, carbonates, 42 concurrent range zones, 146 Condie, K.C., 284 conformities, 192–6 conglomerates, 40, 42, 177 conifers, 259 conodonts, 181, 202, 283 zones, 148, 149 contacts, 114 continental crust, thickening, 216, 217 continental drift, 184 theories, 184 continental environments, 67–81 deposition, 67 ice-dominated, 68–9 transportation, 67 water-dominated, 69–75 wind-dominated, 78–81 continental rifts, 171, 222–3 characteristics, 222–3 occurrence, 222 continental shelves, 81, 85–9 glacially inﬂuenced, 87–8, 90 shallow, 109 storm-dominated, 87, 89 tide-dominated, 86–7, 89 wave-dominated, 87, 89 continental slopes, 81, 88, 89–90, 91 contourites, 89 convolute bedding, 33 convolute lamination, 33 convulsions, worldwide, 187

Conyers, L.B., 281 Copernicus (Moon), 288 coral reefs, 94, 95, 263 distribution, 264 extinction, 272 corals, 50, 51, 53, 92 distribution, 264 fossil, 263 Paleozoic, 263 Coriolis force, 268 coronae, 291 correlation bentonite, 126, 129 local sections, 115 seismic cross-sections, 123 space vs. time, 115 time, 115 see also graphic correlation; physical correlation cosmic rays, 163 Courtillot, Vincent, 267 covalent bonding, 13 Cowie, J.W., 201 Craft, J.H., 99 craters complex, 25 simple, 25 volcanic, 26 see also impact craters cratonic basins, 25, 209, 223 occurrence, 220 subsidence, 220 cratons, 7, 79, 190, 197, 282 Creation, 143 date of, 140–1 Cretaceous, 235 Age, 205 basalts, 170 chalks, 120 clays, 170 climate, 269 dinosaurs, 259 divisions, 201 extraterrestrial impacts, 287 facies, 129, 174 faults, 240 ﬂoods, 246 foraminifera, 154, 157, 169 impact layer, 115 limestones, 18 magnetic reversals, 202 mass extinction, 25 ocean ridges, 244 sea-level changes, 244, 248 supersequences, 198 System, 205 unconformities, 170 vegetation, 260 see also Late Cretaceous; Lower Cretaceous Cretaceous–Cenozoic boundary, 170

Cretaceous–Tertiary boundary, 154, 157, 169 crevasse splays, 62, 73 Crick, R.E., 111 crinoids, 50, 51 Croatia, 54 cross-cutting relationships, principle of, 116, 142 Cruciﬁxion, 140 Cruziana spp., 36 crystallization, dating methods, 163–4 cultivation, 282 Curie point, 155 current–wave bedforms, 29 Cuvier, Georges (1769–1832), 103–4, 129, 187 cyanobacteria, 58, 59, 74, 273 photosynthesis, 267, 271, 272 cycle stratigraphy, 189, 190–2 cycles causes, 195 classiﬁcation, 190 composite, 190–2 deﬁnition, 190 freeze–thaw, 15, 69 icehouse–greenhouse, 190, 265 limestone–marl, 190–2 punctuated aggradational, 7, 196 shallowing-upwards, 196, 197 see also Milankovitch cycles cyclical theories, 116 cyclothems, 190, 192 Cypricardella spp., 133 Dalmatians (dogs), 58–9 Dalrymple, G.B., 162 Dalrymple, R.W., 87 Dana, J.D., 189 Danby, Francis (1793–1861), The Deluge (1840), 188 Darwin, Charles Robert (1809–82), On the Origin of Species by Means of Natural Selection (1859), 141, 143 dating issues Jurassic, 110 volcanic environments, 67 dating methods, 1, 7–8, 9, 115 numerical, 111 relative, 111 sidereal, 278 see also numerical dating methods; radiocarbon dating; radiometric dating Davies, G.F., 216 Davis, J.L., 281 Davis, R.A. Jr., 67 Davis, T.L., 173 Dawson, A.G., 255 de Pater, Imke, 285 Death Valley (USA), 79

326 Index debris avalanches, 25–6, 29, 39, 40 deﬁnition, 27 debris ﬂows, 25, 26, 78 characteristics, 30 ﬂuid, 28, 31 mechanisms, 27 silts, 30 Deccan ﬂood basalts (India), 287 decompaction equation, 179 deconvolution, 121 deep seas, 81 deep-sea carbonates, 96–7 deep-sea deposits, 90–1 deep-sea fans, 90–1, 92 deep-sea plains, 91–3, 127 delta-platform deposits, 83 deltas, 74, 243 ancient, 106 arenites, 42 biofacies, 133 birdfoot, 83, 84, 133 ebb-tidal, 3, 4 fans, 90–1 ice-contact, 245 isostasy, 246 lobes, 108, 116, 117 Nile-type, 83–4 prograding, 83, 89 tide-dominated, 83, 85 wave-dominated, 83–4, 86 see also Mississippi Delta (USA) delta-slope deposits, 83 dendrochronology, 142, 158, 160, 161, 203, 278 calibration, 168, 281 Dennison, J.M., 137 density and gravity, 213 increase, 217 and isostasy, 213–15 reduction, 217 deposition carbonates, 54–5 clastic sediments, 22–3, 24–5 continental environments, 67 limestone, 267 mechanisms, 22–3 sediments, 22–4 traction, 28, 29 wind, 24, 30–2 depositional sequences, 172, 192–6 depositional texture, 38–9, 55 deserts, 15, 69 arid, 81, 260 dunes, 255, 260 environments, 79 formation, 257 polar, 257 sand seas, 80–1 sediments, 82 detrital grains, 24, 271

Deutsch, A., 25 Devonian, 7, 48, 52, 184 epicontinental seas, 97–9 lake deposits, 75 Middle, 133 rifts, 222 sandstones, 105, 177 shales, 132, 133 vegetation, 272 Dewey, J.F., 197, 217, 245 diagenesis, 33, 36, 49, 272 carbonates, 49, 55–9 limestones, 18 peat, 62 and secondary magnetization, 158 diamictites, 25–6, 30, 40 diamond, 25 diaplectic glasses, 25, 26 diastems, 105, 106–11, 150, 151 erosional, 106–9, 110 non-depositional, 106–9 diatoms, 59 Dickin, A.P., 162, 165 Dickson, T., 47 Dietz, R.S., 25 dikes, 33, 40 basalt, 142 dilute water ﬂows, 28 dinosaurs, 170, 259 dip-meter logs, 125, 126 direct solution, 14 directional remnant magnetization (DRM), 155 disconformities, 106, 190, 192 discontinuities, 173 dispersion, 235 disrupted terranes, 235 dissolution, carbonates, 58 distributary channels, 83, 84 divisions, 187 dogs, 58–9 dolomite, 47, 55, 128 precipitation, 60 replacement, 58 dolomitization, 58–9, 94 Donovan, D.T., 75, 244 d’Orbigny, A.D., 5, 187 Dorset (England), 113, 190, 191 Dott, R.H. Jr., 241 downlap, 172 draas, 39 driftwood, 245 drilling, 123, 124, 125 dripstones, 49 Driscoll, N.W., 192 DRM (directional remnant magnetization), 155 dropstones, 87 Dry Valleys (Antarctica), 79 Duff, P. McL. D., 190 Dunbar, C.O., 7, 201

dunes, 33, 39, 78, 114, 246 deserts, 87, 255, 260 eolian, 81 migration, 66, 78 in sand seas, 80 types of, 81 water-laid, 81 Dunham’s classiﬁcation, 55, 56 boundaries, 57 durain, 63 dykes, 217 dynamic corrections, 121 dynamite, 119 Early Cretaceous, magnetic reversals, 202 Earth, 286 age of, 7–8, 140–2 atmosphere, 272 formation theories, 140 geodesy, 209–10 geoidal eustasy, 242–3 heat ﬂow changes, 265–7 orbit, 265, 279 earthquakes, 33, 187 earthworms, 18, 20 East Africa, 222, 268 East Gondwana, 158 Eastern Asia, 243 ecology, 9 ecostratigraphy, 7 Eggleton, R.A., 12, 16 Egypt, 106, 111, 112 Ancient, 160 Eh, 23, 59 Eicher, D.L., 140 Einsele, G., 67, 171 ejecta blankets, 25 impact, 26, 71, 220 volcanic, 26 elephants, 278 Elie de Beaumont, Léonce (1798–1874), 104, 187 Emiliani, Cesare (1922–95), 155 England, 137, 143, 187, 190, 191 ammonites, 110 sandstones, 105, 114 shales, 113 environmental changes, 272 internal vs. external factors, 66 environments ammonites, 144 architectural elements, 67, 70 and carbonate mineralogy, 50 characterization, 12 classiﬁcation, 68, 69 climatic, 15 deﬁnitions, 66–7 marine, 36, 260, 271 mixed, 99–100

Index 327 peculiar, 99–100 processes, 67 recognition, 67 reviews, 67 spatial studies, 9 temporal studies, 9 water, 29 under water, 81–99 see also continental environments; glacial environments; impact environments; volcanic environments Eocene coal seams, 74 conifers, 259 eolian lamination, 38 eolian sands, 79–80, 125 epeiric seas see epicontinental seas ephemeral streams, 78, 81 epiboles, 146 epicontinental seas, 81, 86, 97–9 modeling, 97–9 epifauna, 89 equator, 209, 215 Eratosthenes (c.276–194 BC), 209 Erba, E., 267 ergs, 78 Eriksson, Kenneth A., 284 Ernst, R.E., 285 erosion, 114, 116, 126, 143, 271 base level of, 106–9, 250 differential, 219 local, 105 and sea-level changes, 241 and unconformities, 106 erosional diastems, 106–9, 110 Estonia, 106 estuaries, 83 Eurasia, 15, 81 mammals, 277 Europa (moon), 291 Europe, 147, 184, 187 rock units, 122 eustasy, 9, 173, 241, 253 climatic, 242, 243, 245–6 deﬁnition, 242 geoidal, 242–3 glacial, 245 lacustrine, 245 tectonic, 242, 243–5 and tectonics, 195–6 types of, 242, 243 eustatic correlation charts, 197, 198 eustatic sea-level changes, 130, 178, 192, 196, 241–6 and climate, 268–9 effects, 172, 216, 226 global, 192, 243–4 Quaternary, 278 evaporation, 59, 154, 263 lakes, 78–80

evaporites, 46, 59–62, 284 climatic effects, 260 evaporitic basins, 96 evolution, 100, 143, 144, 171 mammals, 277–8 organisms, 201 plants, 267, 271 of stratigraphic sequences, 208 exinite, 63 exotic terranes, 235 deﬁnition, 209 extinctions, 170, 272 coral reefs, 272 organisms, 201 see also mass extinctions extraclasts, 24, 46 deﬁnition, 54 extraterrestrial impacts, 25–7, 100, 115, 126, 129, 142 effects, 291 evidence, 170 and stratigraphic evolution, 208 Venus, 291 Yucatan, 287 extraterrestrial stratigraphy, 284, 285– 91 issues, 285 Mars, 287, 291 Mercury, 287 Moon, 285–6, 288, 291 and terrestrial stratigraphy, 291 Venus, 287–91 Eyles, C.H., 277 Eyles, N., 277 Eyre, Lake (Australia), 19, 78, 80 facies, 174 concepts of, 5, 129–32, 189 deﬁnition, 127, 129 distributions, 175 impact, 26 migration, 171 models, 283, 284 volcaniclastic, 27 see also biofacies; lithofacies facies analysis, 9, 175 facies changes, 7, 102, 131 in Andean belts, 232 in collision orogens, 234 deﬁnition, 127 in island arcs, 231 in ocean ﬂoors, 225 in passive margins, 229 reef to back-reef, 127–8 in subduction zones, 229 in transform belts, 235 and water depths, 248–53 facies fossils, 144 FAD (ﬁrst appearance datum), 202 Fahraeus, L.E., 149 Fairbridge, R.W., 241

fans deep-sea, 90–1, 92 deep-water, 226 humid, 70 lowstand, 196 outwash, 87 prograding, 92 stream-dominated, 70 subaerial, 90 see also alluvial fans fault planes, Plate 8.1 faulting, 222, 223, 229 transcurrent, 185 see also strike-slip faulting faults, 125, 187, 229, 240 half-graben, Plate 8.5 transcurrent, 222 fauna, provincial linking, 235 faunal distribution, and tectonics, 147 Faunal Succession, Law of, 129 faunas, 94, 189, 222, 233, 263 Cambrian, 275 catastrophic creations and extinctions, 103–4 epifauna, 89 facies changes, 129 fecal pellets, 21, 24, 50, 54 peloid formation, 94 feldspar, 42 feldspathic greywackes, 42 fence diagrams, 171–2 ferns, 63 ferrasols, 18 ferric iron, 15 ferric oxide, 59, 260 ferroan carbonates, 47 ferromagnesian grains, 36 ferrous iron, 15 oxidation, 271 ferrous oxides, 15 ﬁreclays, 46 ﬁrst appearance datum (FAD), 202 Fischer, A.G., 190 ﬁsh, 51, 75, 170 fossil, 148 Fisher, Richard V., 67 ﬁssion track dating method, 162, 168–70, 181, 183 ﬁssure volcanoes, 26 Fitzpatrick, E.A., 12 ﬂaser bedding, 29, 35 ﬂash ﬂoods, 78 ﬂexural rigidity, of lithosphere, 215, 216 ﬂoating equilibrium, 213–15 ﬂoating textures, 38–9 ﬂood basalts, 267, 286 Deccan, 287 ﬂoodplains, 70–2, 73, 106 ﬂoods, 28, 74, 187, 246 Mars, 287

328 Index ﬂoras, 63 Cenozoic, 262 Florence (Italy), 116 Flotz-Geburge, 187 ﬂowstones, 49 Flugel, E., 47, 54, 67 ﬂuid debris ﬂows, 28, 31 ﬂuid ﬂows, 22, 24–5, 40 ﬂuorapatites, 65 Fly Delta (New Guinea), 85 folds, 125 Folk’s classiﬁcation, 55 foraminifera, 53, 93, 154 benthic, 155, 169, 265 planktonic, 155, 265 zones, 145 foraminiferal oozes, 50, 223 forebulges, 215, 246 collapse, 246 migrating, 217 foreland basins, 232, 233, 234 foresets, 74 formation density logs, 124 formations, 122–3, 172, 200 concepts of, 111–14 subdivisions, 114 forts, 280 fossil animals, 263 fossil corals, 263 fossil plants, 262 fossil space, concept of, 144 fossil zones, 145–52, 199 characteristics, 146–8 correlation issues, 146 deﬁnition, 145 extent, 148–9 graphic correlation, 149–52 and lithofacies, 145, 150 Middle Jurassic, 199 types of, 145–6 see also zone fossils fossiliferous rocks, 187 fossils, 53, 54, 104–5, 111 accumulation rates, 149–50 benthic, 7, 230 biofacies, 133–4 cementation, 36 characteristic, 118 common, 143 in dating methods, 115–16, 142, 143–52, 170 early studies, 7 and facies, 127, 128–9, 130–1, 132 facies, 144 index, 146 marine, 3, 48, 75 Paleozoic, 100 physical correlation, 115 in reefs, 96 and relative ages, 143–52 and rock units, 187

shelly, 283–4 and strata, 4, 5, 7, 116 taxonomy, 143–4 and time units, 103 trace, 33–6, 41, 190, 272 see also zone fossils Frakes, L.A., 255 framework textures, 39, 40, 43 France, 130, 151, 187, 199, 209, 241 facies changes, 7 Fraser, G.S., 67 free air correction, 210–11 freeze–thaw cycles, 15, 69 French, B.M., 25 freshwater, 65, 263 evaporation, 59 isotopic composition, 263 oxygen isotopes, 154 Frey, R.W., 21 Friedman, G.M., 47 Fuller, J.G.C.M., 143 fungal hyphae, 168 fungi, 18 fusain, 63 Galilei, Galileo (1564–1642), 210 Galloway, W.E., 192 gals, 210 gamma radiation, 124, 125 Ganges valley (India), 70 Gang-Geburge, 187 gas, 24, 174, 184 formation, 181 gas ﬁelds, 1–3 gastropods, 54, 75 Gauss–Matyuma boundary, 278 Gauss–Matyuma paleomagnetic reversal, 155 Geer, Gerard Jacob de (1858–1943), 160 Geldart, L.P., 118 genetic sequence stratigraphy, 192–7 assumptions, 192 sequence generation, 192–6 geochemistry and relative ages, 152–3 volcanic ashes, 126–7, 153 geochronology, 103, 199–200, 203–5, 283 Quaternary, 278 geodesy, 209–10 deﬁnition, 209 geoidal eustasy, 242–3 geoids, 210 anomalies, 242–3 oceanic, 241 geologic history, interpretation, 9 geologic time, scale of, 140 Geological Society of America, 7 geology development, 3–8 golden age of, 104

geomagnetic ﬁelds, 157 geophones, 119–120 geophysics, 9 geopotentials, 210 George (hurricane), 88 George, T.N., 192 geothermal gradients, 170 Germany, 145, 148, 151, 187 craters, 26, 220 Gibraltar, Straits of, 243 Gilluly, J., 189 glacial deposits, 81, 283 marine, 70 stratigraphy, 276–7 glacial environments, 15 oxygen isotopes, 155 sub-environments, 68–9 types of, 73 glacial eustasy, 245 glacial stages, 277 glacial valleys, 111 glaciations, 261, 265, 267 glaciers, 36, 37, 82, 86, 263–5 cycles, 268, 277 grounded, 196 isotopic composition, 154 kinetics, 22, 25, 27 Mars, 287 mechanisms, 30 occurrence, 68 Quaternary, 246 types of, 73 glasses, 168 diaplectic, 25, 26 glauconite, 65, 262 Glen Canyon Group (North America), 114 gleyed soils, 21 global sea level, 196–7 Global Stratotype Section and Point (GSSP), 201 global warming, mechanisms, 269 Glooschenko, W.A., 62, 262 gneisses, 42 God, 143 goethite, 65 Gohau, G., 3 golden spikes, 201 Gondwanaland, 158, 180, 244, 260, 261 Goodman, D., 281 Goodwin, P.W., 196 Gould, S.J., 130, 140, 186 Grabau, Amadeus William (1870–1946), 7, 187, 190 graded beds, 28, 190 grain compaction, carbonates, 58 grain composition, 36 grain ﬂows, 28, 31, 39 grain roundness, 38, 43 grain shape, 38

Index 329 grain size, 36–8, 53 clastic sediments, 42 distribution, 38 sediments, 81 variable, 24 grains coated, 54 detrital, 24, 271 ferromagnesian, 36 non-skeletal, 54 quartz, 23, 30, 38 skeletal, 53–4 grainstones, 55, 57, 58 El Gran Desierto (Mexico), 81 Grand Canyon (USA), 116, 118, 132–3 granite domes, 282 granites, 20, 38, 42 chemical decomposition, 14 granodiorite–granite intrusions, 232 grapestones, 24, 54 graphic correlation fossil zones, 149–52 in magnetostratigraphy, 158 graptolites, 7, 35 zones, 147, 148, 149 gravels, 22, 25, 27, 78 beach, 131 cross-bedded, 70 graded, 96 gravity, 209, 210–13, 241 effects on sediments, 22 units, 210 gravity anomalies, 211–13 local, 213 gravity measurements, 210 corrections, 210–13 Great Salt Lake (USA), 79 Greece, Ancient, 209, 283 green algae, 50, 51, 54 greenhouse gases changes, 267–8 determinants, 267 Greenland, 154, 184 Greensand–Gault–Chalk succession, 129 greenstone belts, 282, 284 Gressly, A., 7, 130–1, 187, 190 greywackes, 36 lithic, 42, 45 Grieve, R.A.F., 25 ground-penetrating radar, 281 groundwaters, 16, 24, 36 evaporation, 59 groups, 114, 200 GSSP (Global Stratotype Section and Point), 201 Guinea, Gulf of, Plate 8.3 Gulf Stream, 268, 269 Gurnis, M., 243 gymnosperms, 63 gypsum, 59, 60, 62, 94, 260

Hadley circulation, 255–7 half-lives see radioactive decay rates Halfway Formation (Canada), 2, 3 Halimeda spp. (green algae), 51 halite, 59–60, 260 halite pseudomorphs, 33 Hall, James (1811–98), 7 Hallam, A., 129, 190, 241, 244, 246, 248 Halley, Sir Edmond (1656–1742), 141 Hamblin, W.K., 285 Hamilton, W.B., 240, 284 Haq, B.U., 244 Harder, H., 271 hardgrounds, 58 Hardie, L.A., 67 hardwood, density, 214 Harland, W.B., 199, 201 Hart, B.S., 174 Hatteras abyssal plain (Atlantic), 93 Hauterivian Age, 205 Hauterivian Stage, 205 Hawaii, 159, 216, 223 Hay, W.W., 210, 245, 268 Head, J.W., 291 heat budgets, 265 distribution, 265 heat ﬂow, 276, 284 changes, 265–7 heliacal risings, 160 Helwig, J., 235 hematite, 65 Hettangian, 247 marine sedimentation, 248 highlands, lunar, 288 high-magnesium calcite, 47, 50, 54, 57–8 precipitation, 60 highstand systems tract, 196 high-strontium calcite, 54 Hilgen, F.J., 279 Himalayas, 15, 29, 70, 180, 218 collision orogens, 232 uplifts, 218 Hindeodus parvus (conodont), 202 Hindu culture, 140 Hinnov, L.A., 192 history, concept of, 140–1 Hjulström’s diagram, 23 Holland, 52 Holland, C.H., 153, 199 Homer (8th cent. BC), Iliad, 248 Hongfu, Yin, 202 horizontal stacking, 121 hornblende, age spectra, 167 House, M., 192 Howe, R.C., 208 Hubbard, R.J., 197, 243 Hudson Bay (Canada), 81, 97, 99 impact basin, 220, 222, 223 Huff, W.D., 127, 201 huminite, 63 hummocky cross-stratiﬁcation, 29, 34

humus, 18–20 Hurricane George (1998), 88 hurricane waves, 29 Hutton, James (1726–97), 105, 129, 141, 187, 190 catastrophism theories, 130 cyclical theories, 116, 141 Hyde, R.S., 284 hydration, 14, 245–6 mechanisms, 15 hydrocarbons, 174, 182, 184 formation, 181 see also gas; oil hydrodynamic regimes, 91 hydrogen ions, 15 concentrations, 124–5 hydrolysis, mechanisms, 14 hydrophones, 119–120 hydroxyapatites, 65 hyperconcentrated ﬂows, mechanisms, 27–8 hypercyclothems, 190 hypsometric curves, 210, 211 hypsometric integrals, 210 hypsometry, 210, 243, 246, 247 changes, 248, 265, 268–9 deﬁnition, 209 factors affecting, 210 Iapetus Ocean, 235, 274 ice density, 214 isotopic composition, 154–5, 156 ice ages, 154, 265, 267, 276 global, 283 ice caps, 261, 268 ﬂexural loading, 215 formation, 242, 255 melting, 215, 242 see also poles ice cores, 154 ice ﬂows, 22, 25, 30, 67 ice sheets, 192, 246, 277 melting, 278 icebergs, 30, 214 density, 214–15 ice-contact deposits, 68, 69 ice-dominated systems, 68–9 icehouse–greenhouse cycles, 190, 265 Iceland, 184, 258 glaciers, 268 ichnofacies, 36 Idaho (USA), 64 igneous processes, 276 igneous rocks, 38, 106, 158, 200 crystallization, 13 weathering, 15, 269 illite, 45 Imbrie, J., 190, 265 Imbrie, K.P., 190, 265 immature sediments, 24

330 Index impact craters, 25, 170, 220, 285 complex, 220 densities, 285 Mercury, 289 Moon, 285 simple, 220 Venus, 291 see also extraterrestrial impacts impact environments, 67 recognition, 67 sedimentary sequences, 71 impact pressures, 25 incised streams, 73 incised valley ﬁlls, 196 index fossils, 146 India, 18, 29, 160, 268, 287 fans, 70, 76 Indian Ocean, 268 Indonesia, 231 Indus, River, 91 inertinite, 63 Ingersoll R.V., 221 initial horizontality, principle of, 116 inner planets, 285, 286 instrument logs see wireline logs interglacial stages, 277 International Stratigraphic Guide (1994), 199, 201 interstadials, 277 interval zones, 146 intraclasts, 24, 46 deﬁnition, 54 ions, 14, 15, 20, 47, 125 calcium, 49 concentration, 50 transportation, 22 and weathering, 13 Ireland, 141 iridium, 170 Irish Sea, 179, 181–4 iron, 58 colloidal, 65 ferric, 15 ferrous, 15, 271 iron compounds oxidation, 15 reduction, 15 iron ores, oolitic, 65 iron oxide, 20 ironstones, 59, 65 banded, 100, 271, 283, 284 oolitic, 100 island arcs, 228, 230–1, 232 island-arc volcaniclastics, 99, 231 isochronous surfaces, 192 isochrons, 165, 170 isopach maps, 139, 174, 174–5 isostages, 203 isostasy, 9, 213–18, 241, 246, 253 deﬁnition, 209 isotope ratios, 47–8, 50, 163–4

ancient, 263 changes, 153–4 isotope stratigraphy, 203 isotopes biochemical precipitation, 47 breakdown, 168 carbonates, 47–8 chemical precipitation, 47 radioactive decay rates, 162–3, 164–8 unstable, 162 see also carbon isotopes; oxygen isotopes; stable isotopes; uranium isotopes isotopic exchange, 168 Italy, 116, 187, 266 Apennines, 219, Plate 8.4 stratigraphy, 116, 117, 142, 143 James Bay (Canada), 222 James, N.P., 50, 67, 228, 272 Japan, 119, 147, 230, 231 Japanese Islands, 232 Johnson, M.E., 189 Joly, John (1857–1933), 141 Jones, E.J.W., 244 Jordan, Plate 14.1 Jura mountains, 5–7, 131 Jura Quartzite, 31 Jurassic, 5–7, 52, 187, 240 ammonites, 110 dating issues, 110 ﬂuid debris ﬂows, 31 fossil zones, 145 limestone deposition, 267 magnetic reversals, 202 sandstones, 114 timescales, 192 see also Lower Jurassic; Middle Jurassic; Upper Jurassic Jurassic–Cretaceous passive margins, 197 kalpas, 140 Kansas (USA), 193 kaolinite, 45, 46 katabatic winds, 45, 81 Kauffman, E.G., 201 Kay, G.M., 190, 200 Kayenta Formation (USA), 123 Keller, E.A., 210 Kelvin, William Thomson, 1st Baron (1824–1907), 7–8, 141–2, 189 Kendall, A.C., 60 Kentucky (USA), 108 Kimmeridgian, 147 King, Philip B., 127 Klein, G. de V., 192 Knapp, B., 12 Kocurek, G.A., 78 Kolata, D.R., 220 komatiitic basalts, 284 Kon Tiki Expedition, 214 Kongur, Mount (China), 82

Kosi River fan (India), 70, 76 Kuenen, H., 246 Kunlun range (China), 82 Kyrgystan, 111 lacustrine eustasy, 245 lag deposits, 78, 80 lagoons, 94, 95 Lahul (India), 29 Laityau Petchabun-Khiong Marui fault (China), 240 lakes, 67 carbonate sediments, 78 characteristics, 78 chemical changes, 152 evaporation, 78–80 evaporites, 60 glacial, 68, 69, 160 large, 73, 74–5 oxbow, 66 rift, 222 varved clays, 262 see also saline lakes lamination convolute, 33 eolian, 38 land distribution, changes, 268–9 landslides, 187 Lapworth, Charles (1842–1920), 105 Larson, R.L., 267 Late Cretaceous, 63, 265, 268 impact basins, 220 Late Precambrian, 106 polar wanderings, 158 lateral changes, 127–34, 281–2 lateral ﬂow, 217, 218 lateral tracing, of strata, 118 laterites, 21 lattice structures, 162, 165 Laurentia, 158 lava beds, 200 lavas, 26, 67, 157 basalt, 170, 223 pillow, 159 Venus, 291 leaching, 21 lead isotopes, 163 leaf anatomies, 262, 263 Leeder, M.R., 22 Lehmann, Johann Gottlob (1719–67), 187 Leighton, M.W., 220 lenses, 114, 117, 189, 192, 200 meltrock, 25, 220 stratigraphy, 277 lenticular bedding, 35 levees, 66, 73 Lias, 192 Blue, 191 Lower, 113 limestone deposition, and greenhouse gases, 267

Index 331 limestone–marl cycles, 190–2 limestones, 48, 54, 60, 93, 128, 220 ammonites, 151 caves, 280 chert nodules, 59 clastic, 55, 94 clayey, 190 compaction, 179 diagenesis, 18 erosion, 46 formation, 55–6, 223, 272 formations, 111, 114 interbedded, 114, 115, 190 marine, 192 mineralogy, 53 Paleozoic, 259, 272 Precambrian, 272 reef, 131 relative ages, 143 shelly, 64 Silurian, 105 Upper Jurassic, 130 weathering, 15 limonite, 64, 65 lineage zones, 146 Lingulida (brachiopods), 144 liquids, 24 Lissauer, J.J., 285 Lister, A.M., 278 lithic arenites, 42, 45 lithic greywackes, 42, 45 lithiﬁcation, carbonates, 58 lithodemic units, 200 lithofacies, 131, 132–3, 135, 172 changes, 222 and fossil zones, 145, 150 in ocean ﬂoors, 223–5 lithological similarity, 122–7 lithosphere, 214–15, 228, 230 ﬂexural loading, 215–16 ﬂexural rigidity, 215, 216 see also oceanic lithosphere lithostratigraphic classiﬁcation, 103, 104 lithostratigraphic units, 103, 111, 200–1 lithostratigraphy, 187, 200–1 approaches, 4–5 rock units, 201 see also rock units Little Ice Age, 265 loading, 33 local sections, 103–5 correlation, 115 subdivision, 111–14 loess, 46, 81 Logan, B.W., 60 Lophelia spp. (corals), 92 Lowe, D.R., 284 Lower Cretaceous, 135, 153, 197 orbital cycles, 266

Lower Jurassic, 64, 147, 148, 190, 191 ammonites, 254 marine sedimentation, 248 Lower Lias (England), 113 Lower Middle Natu Formation (China), 175 Lower Paleozoic, 7, 126 divisions, 104 Lower Triassic, 179 low-magnesium calcite, 47, 58 precipitation, 60 Lowman, P.D. Jr., 285 lowstand fans, 196 lowstand system tracts, 196 Ludvigsen, R., 145, 149 Luthi, S.M., 124 Lyell, Sir Charles (1797–1875), 7, 130, 187, 189 Principles of Geology (1833), 141 McArthur, J.M., 154 McCabe, P.J., 62 McKenzie, D.P., 177, 291 Mackenzie, F.T., 47, 265 Mackenzie, J.A., 265 McLane, M., 22 McQuillan, R., 122 macrocycles, 190 Mae Ping fault (China), 240 magmas, 165 basaltic, 217 magmatic arcs, 217, 223, 228, 231, 232, 244 magmatism, 208, 222 island arcs, 230 magnacycles, 190 magnesite, 47 magnesium, 153 magnesium carbonate, 58 magnesium cations, 47, 50 magnesium poisoning, 50 magnesium salts, 60 magnetic anomalies, 268 magnetic ﬁelds, 155–6 magnetic materials, 155–6 magnetic reversals, 103, 115, 142, 157–8, 202 timescales, 142, 157, 159, 160 see also paleomagnetic reversals magnetite, 155 magnetization, secondary, 158 magnetostratigraphic polarity zones, 202 magnetostratigraphy, 157–8, 160, 202 advantages, 157–8 disadvantages, 284 graphic correlation, 158 limitations, 158 Maine (USA), 246 mammals, 170 evolution, 277–8 manganese, 58

mantle, 164, 213, 242, 284 degassing, 269 rheology, 215 mantle–crust columns, 178 mantle plumes, 220 mapping, 134–9 early, 104–5 units, 122–3 maps, 8, 9 amplitude, Plate 8.2 features, 134 geological, 102, 111, 134–9 isopach, 139, 174, 174–5 limitations, 137–9 outcrop, 134–5, 137 paleogeographic, 184–5, 260 paleotectonic, 184–5 Mare Imbrium (Moon), 288 maria, 285 marine clays, 246 marine environments, 36, 260, 271 marine oozes, 47 marine sedimentation, 248 marine sediments, 81, 155, 228, 231 Markov chain analysis, 190 Mars, 208, 241, 286 atmosphere, 272, 287 rifts, 287 stratigraphy, 287, 291 surveys, 285 volcanoes, 287 marshes, 73, 84 Martini, I.P., 62, 196, 246, 262, 267 mass extinctions, 25 catastrophic, 103, 104, 129 evidence, 170 mass ﬂows, 22, 24–5, 27–9, 40, 67, 82 deposits, 67, 69, 78, 89, 260 in ice-dominated systems, 68 Mars, 287, 290 in reefs, 96 mature sediments, 24 maturity concept, 24 maximum ﬂooding surface, 196 Mayas, 140 meandering channels, 66, Plate 8.2 modeling, Plate 8.3 meandering streams, 69–70, 72, 177 characteristics, 73 Mediterranean Sea, 111, 243 megacycles, 190 megacyclothems, 190 Meishan section (China), 202 Mel’nyuk, Iu. P., 100 Melosh, H.J., 25, 67, 220, 284 meltrock lenses, 25, 220 meltwater streams, 68–9 Melvin, J.L., 67 members, 114, 200

332 Index Mercury, 286 impact craters, 287, 289 stratigraphy, 287 Merk, G., 7 mesocycles, 190 Mesoproterozoic, 283 Mesosaurus spp., 1 Mesozoic, 51, 240, 255 bioturbation, 272 glaciation, 26 plankton, 272 sea-level changes, 252 sequence boundaries, 197 subsidence, 179 terranes, 240 Mesozoic–Cenozoic boundary, 170, 240 metal ions, 47 metal sulﬁdes, 271 metamorphic rocks, 38, 106, 200, 284 weathering, 15, 269 metamorphic terranes, 235 metamorphism, 165, 189, 223, 229, 276, 284 hydrothermal, 223, 225 regional, 231, 232, 234 and secondary magnetization, 158 welding, 235, 238 meteorites, 100, 129, 170, 284 bombardment, 285 lunar ﬂux, 286, 287 Mexico, 81, 170, 220, 221 Mexico, Gulf of, 88, Plate 8.2 Miall, A.D., 67, 70, 171, 190, 197, 221 Miall, C.E., 190 Michigan basin (USA), 220 micrite, 59 micritization, 58 Mid-Atlantic Ridge, 159, 213 Middle Ages, 265 Middle Devonian, 133 Middle Jurassic, 157 fossil zones, 199 Middle Ordovician, 97 trilobites, 147 Middle Oxfordian, 131 Middle Triassic, 1, 6, 179 Middleton, G.V., 11 mid-oceanic rifts, 222, 223 migration (seismic correction), 122 Milankovitch cycles, 190, 192, 265, 266 in astrochronology, 279 correlations, 192 Milky Way, 265 Mill, John Stuart (1806–73), A System of Logic (1843), 186 milligals, 210 Milliman, J.D., 47 minerals crystallization, 163–4 formation, 14 lattice structure, 162, 165

potassium-bearing, 165 radioactive, 162 uranium-bearing, 168 mining, 186–7 Miocene, 63, 153, 240, 243, 279 Mississippi Basin (USA), 174 Mississippi Delta (USA), 136 aeronautical charts, 138 barrier complex, 87, 88 birdfoot delta, 84, 133 disconformities, 192 lobes, 108, 116, 117 Mississippi River (USA), 133–4, 197, 210 embayment, 263 river terraces, 111 Mitchum, R.M. Jr., 192 moder humus, 20 mollusks, 50, 51, 94, 133 Cenozoic, 263 Monger, J.W.H., 235 monsoons, 268, 269 montmorillonite, 45 Moon, 286 impact craters, 220, 285 meteorite ﬂux, 286, 287 stratigraphy, 285–6, 288, 291 surface types, 285–6 surveys, 285 moons, stratigraphy, 291 Moore, P.D., 62 mor humus, 20 moraines, submarine, 87 Morrison, J.O., 47 Mount St Helens (USA), 128, 216 mountain belts, 235 analyses, 209 see also orogens mountain building, 15, 217 worldwide, 188, 189 Muav Limestone (USA), 132 Mucci, A., 47 mudcracks, 33 mudﬂows, 23 muds, 57, 83, 95, 96, 125 carbonate, 54, 131 mudstones, 28, 30, 42–5, 46, 93 calcareous, 190 formation, 74 mull humus, 18 Murchison, Sir Roderick Impey (1792–1871), 4–5, 7, 104–5 Murphy, M.A., 145 muscovite, 165 Naeser, C.W., 168 Naeser, N.D., 168 Natron, Lake (Tanzania), 79 natural gamma ray logs, 125, 126 Navajo Sandstone (North America), 114, 123 near-surface physical changes, 32, 33–6

Nelson, C.H., 28 Nelson, Clifford M., 130 neodymium, radioactive decay rates, 164 neomorphism, carbonates, 58 Neoproterozoic, 99, 283–4 Neotethys ocean, 180 Neukum, G., 287 neutron logs, 124–5 neutrons, 163, 168 New Guinea, 85, 243 New York (USA), 143 New York State (USA), 7, 132, 133 New Zealand, 50 Newton, Sir Isaac (1642–1727), 209 Niger delta, 86 Nile canyon, 111, 112 Nile Delta, 86, 136 Nile, River, 83–4, 106, 112 Nimmo, F., 291 Nisbet, E.G., 284 nitrogen isotopes, half-lives, 163 Noah’s Flood, 116, 187 nonconformities, 106 non-depositional diastems, 106–9 Nordlinger church, 220 North Africa, climate, 255 North America, 97–9, 149, 228 amalgamation, 235 bentonite, 126 carbonates, 122 coals, 74 collision orogens, 232 conodont zones, 149 cyclothems, 190, 192, 194 deserts, 15 divisions, 201 fossil animals, 263 glacial deposits, 277 leaf anatomies, 263 mountains, 15, 184 paleogeographic maps, 184 paleosols, 81 passive margins, 229 rock units, 122–3 sandstones, 114, 123 sea-level changes, 246 sedimentary lenses, 189 stratigraphy, 7, 189 see also Canada; Pennsylvanian (North America); United States of America (USA) North American Commission on Stratigraphic Nomenclature, 199, 199 North American Cordillera, 237 North American Stratigraphic Code (1983), 199, 201 North Atlantic, 185, 235, 257 Gulf Stream, 269 North Atlantic Deep Water Current, 269 North Paciﬁc, 257, 258

Index 333 North Pole (Australia), 273 North Sea, 120 North Wales, 5, 104–5 Norway, 92 numerical dating methods, 142, 158–70, 276 calibration, 170, 283 imprecise, 162–70 precise, 158–61, 170 see also radiometric dating numerical time units, 103, 203–5 Nunn, P.D., 242 nutrients ﬂux, 271 turnover, 97 Oberbeck, V.R., 26 obsidian, 168 ocean ﬂoors, 223–6 ocean ridges, 244 oceanic lithosphere, 217, 218 magnetic reversals, 157 oceans, 89 age estimates, 141 areas, 243–4 chemical changes, 152–3, 154 circulation, 257, 258, 268 dating issues, 7 deep seas, 81 depths, 244–5 development, 275 hydrodynamic regimes, 91 Iapetus, 235, 274 Indian, 268 isotopic composition, 154–5 Neotethys, 180 organism-induced changes, 271 stable isotopes, 263 stratiﬁcation, 155 subduction zones, 228–30 temperature, 257, 263–5 upwelling currents, 262 see also Atlantic Ocean; epicontinental seas; Paciﬁc Ocean; Universal Ocean Odin, G.S., 248 oil, 174 formation, 181 oil ﬁelds, 1–3 Oligocene, 232, 240 climate, 265 oncoids, 54, 74 onlap, 172 Onny River (England), 137 Ontario (Canada) Ordovician, 35, 39 Quaternary, 32, 40, 277 Ontong Java plateau (Paciﬁc), 267 ooids, 50, 51, 57, 64 formation, 47, 54 ooliths, 24 oolitic bars, 94

oolitic ironstones, 100 opening and closing oceans model, 221 Oppel, A., 5, 130, 131, 145–6, 187 Oppel zones, 146 Orcadian Basin (Scotland), 75 Ordovician, 7, 35, 105, 137, 235 bentonite correlations, 129 carbonates, 122 climate, 267 ﬂoods, 246 glaciations, 265 graptolite zones, 149 ice caps, 261 sandstones, 32 sea-level changes, 247, 248 Upper, 39 varved clays, 161 see also Middle Ordovician Oreskes, N., 184 organic communities, 50–1, 52 organic matter, 36, 73 decomposition, 18–20 transformation, 16 organisms biogeography, 149 bioturbation, 272 calcium carbonate mineralogy, 47, 48, 50–1 deposits, 83 distribution, 131–2 ecological factors, 51 effects: on beds, 272; on bioclastic sediments, 271–2; on stratigraphy, 271 evolution, 201 extinctions, 201 nektonic, 225, 226 nutrient turnover, 97 organic carbon, 166 pelagic, 225 planktonic, 226 and relative sea-level changes, 253 salinity responses, 134 secretions, 53 skeletons, 53, 54, 283 taxonomy, 53 Orientale Basin (Moon), 288 original continuity, principle of, 116 orogenic belts, 235 terrane analysis, 235–40 orogenies, 184, 189 orogens, 91, 174 collision, 232–4, 235 see also mountain belts orthobreccias, 40 orthoclase, chemical decomposition, 14 orthoconglomerates, 40 Orton, G.J., 67 ostracods, 75, 94 outcrop maps, 134–5, 137

outcrops, 121, 122–3, 143, 171 sequence stratigraphy, 172 outer planets, moons, 291 outwash fans, 87 overlap sequences, 235, 238 owls, 277 oxbow lakes, 66 Oxfordian, 110, 131, 147, 247 marine sedimentation, 248 Middle, 131 Upper, 131 oxidation, 14, 15, 65 mechanisms, 15 oxides, 65 oxygen atmospheric, 271 increase, 267–8 variations, 267 oxygen ions, 15, 47 oxygen isotopes, 48, 49, 154–5, 202 climatic effects, 263–5 marine, 279 variations, 155, 264, 278 Ozarkian system, 7 Ozima, M., 162 Paciﬁc Northwest, 263 Paciﬁc Ocean, 147, 231, 243, 267 Central, 157 North, 257, 258 plateaus, 244 trilobites, 274 packstones, 55 PACs (punctuated aggradational cycles), 7, 196 paleobathymetry, 177, 179, 181 Paleocene, unconformities, 170 paleocurrents, 174, 175, 177 paleogeographic maps, 184–5, 260 paleolatitudes, 158 rock units, 156 paleomagnetic poles, 156–7 paleomagnetic reversals, 153, 170, 279 Gauss–Matyuma, 155 paleomagnetism, 262, 268, 275 advantages, 156 pole positions, 156–7 and relative ages, 155–8 Paleoproterozoic, 283 paleosols, 81, 260 paleotectonic maps, 184–5 paleotemperatures, 48, 183 paleothermometry, 180–4 Paleozoic, 7, 184, 283 bioturbation, 272 clastics, 184 corals, 263 deserts, 260 faunas, 189 ﬂoras, 63 fossils, 100

334 Index Paleozoic (cont.) hypsometry, 210 limestones, 53, 56, 259, 272 oceans, 275 paleotemperatures, 48 polar migrations, 261 terranes, 239, 240 see also Lower Paleozoic Pappalardo, R.T., 291 parabreccias, 40 paraconformities, 106 paraconglomerates, 40 parasequence sets, 196, 197 parasequences, 196, 197 Paris (France), 209 Paris Basin (France), 241 Parrish, J.Y., 260 passive margins, 7, 226–8, 229 cross-sections, 228 sediment ﬁlls, 226–8 Patrick, R.R., 208 PDB (Pee Dee Formation Belemnite), 48 peats, 62–3, 67, 73, 246, 267, 268 formation, 20, 73 Recent, 262 pebbles, 38, 40, 69, 175 pedostratigraphic units, 201 Pee Dee Formation Belemnite (PDB), 48 pelagic sediments, occurrence, 91 peloids, 51, 54, 94 Peltier, W.R., 215, 242 Pemberton, S.G., 36 pendulums, 209 Pennsylvania (USA), 177, 193 Pennsylvanian (North America), 73–4, 107 cyclothems, 190, 192 Permian, 5, 64, 180, 184 breccias, 175, 176 coral reefs, 272 cyclothems, 192 facies changes, 127–8 ice ages, 267 ice caps, 261 ocean ridges, 244 plate tectonics, 244 sandstones, 126 Upper, 185, 203, 260 varved clays, 161 Permo-Carboniferous, 183 cyclothems, 193 Persian Gulf, 54, 95 Petra (Jordan), Plate 14.1 pH, 23, 50, 59 Phanerozoic, 65, 75, 129, 203, 267 stratigraphic system, 204, 205 stratigraphy, 276, 283 Phillips, T.L., 63 phosphates, 58, 65, 262 phosphorites, 64 photic zone, 272

photosynthesis, 15, 46–7, 48, 271, 272 and greenhouse gases, 267 physical correlation, 115–27 vs. time correlation, 115 physical weathering mechanisms, 13 rates, 15 phytoliths, 14, 18 phytoplankton blooms, 160 Pictonia spp. (ammonites), 147 Pierce, R.W., 150 Piggott, S., 281 pigs, 278 Pilbara Craton (Australia), 282 pillow lavas, 159 pines, 160, 259 Pinter, N., 210 pisoids, 54 Pitman, W.C., 197, 244, 245, 246 pits, 281 planets, 140 inner, 285, 286 moons, 291 rotational stability, 215 surveys, 285 see also Earth; Mars; Mercury; Venus plankton, 96–7, 155, 262, 265, 272 plants decomposition, 16 evolution, 267–8, 271 fossil, 262 and soil formation, 18 see also ﬂoras; vegetation planula zone, 131 plate tectonics, 208, 261, 262, 267, 283 Archean, 284 and basin development, 220–35 effects, 209: on climate, 268–9 models, 221–2 paleomagnetic determination, 185 Permian, 244 Plato (c.428–c.348 BC), 249 playas, 78 formation, 78–80 Pleistocene, 215 pluton stitching, 235, 238 plutonism, 234 point bars, 66, 73 polar bears, 133 polar wandering curves, 157, 158 poles, 209 isostatic depression, 215 migration, 261 models, 268 paleomagnetic, 156–7 see also ice caps pollen, 63 polysaccharides, 18, 20 Popigai crater (Siberia), 26

porewaters, 24, 33, 36, 50, 235 carbonate-rich, 272 isotopic exchange, 168 saline, 124 salt removal, 141 supersaturated, 58 porosity, evaluation, 124–5 porosity logs, 124 Portage Glacier (Alaska), 36 Porter, R., 4 post-depositional modiﬁcations, 32–6 potassium, 125, 164 radioactive decay rates, 162, 165–6 potassium–argon dating method, 163, 165–6, 170, 279 potassium salts, 60 potassium-bearing minerals, 165 pottery, 282 Pratt isostasy, 214 Precambrian, 31, 40, 65, 164 bioturbation, 272 divisions, 205 fossils, 284 limestones, 272 rock units, 200 sediments, 171 varved clays, 161 see also Late Precambrian precipitates, 114 formation, 15, 20 semi-arid, 94 precipitation calcite, 47, 50 carbonates, 14, 36, 46, 57, 154 cements, 39 freshwater, 49 isotopes, 47 saltwater, 50 silica, 59, 271 precise numerical ages, 158–61 pressure solution, 58 Primary, 187 Prockter, L.M., 291 Productella spp., 133 Proterozoic, 267, 271 stratigraphy, 276, 281, 283–4 Prothero, D.R., 22, 67 provenance, sediments, 175 provenance linking, 235 provincial linking, 235 pteropod oozes, 50 Puchezh-Katunki crater (Russia), 26 punctuated aggradational cycles (PACs), 7, 196 pyrite, 65, 272 pyroclastics, 26 quarries, 142, 144 quartz, 13, 20, 24, 25 quartz arenites, 42, 45 quartz grains, 23, 30, 38

Index 335 quartz sands, 43, 81 quartzites, 271 Quaternary, 32, 37, 40, 45, 137 astrochronology, 278–9 biostratigraphy, 277–8 chronostratigraphy, 277, 278 clays, 106 climatic change, 111, 278 deposits, 190 eustatic sea-level changes, 278 foraminifera, 265 geochronology, 278 glacial deposits, 81 glaciations, 267 glaciers, 246 ice ages, 154, 265 isotopic variations, 263–5 magnetic reversals, 202 Milankovitch cycles, 279 river terraces, 111 sea-level changes, 246 shorelines, 250 stratigraphy, 276–9 varved clays, 161 Quenstedt, F.A., 145 querns, 282 Quinag (Scotland), 106 Rabaul volcano (New Britain Island), 128 radar, ground-penetrating, 281 radiation, solar, 265 radioactive damage, 168 radioactive decay rates, 162, 163, 164–8, 278 logarithmic, 162 measurement, 162–3 radioactive minerals, 162 radioactivity, discovery, 8, 141–2, 189 radiocarbon dating, 166–8, 216, 278 calibration, 160, 168, 170, 281 errors, 168 radiolarian cherts, 93, 223 radiolarians, 59 radiometric dating, 9, 103, 142, 158, 162–70, 203 accuracy, 284 argon–argon method, 166, 170, 279 assumptions, 162 calibration, 157 carbon-14 method, 142, 163, 166–8 characteristics, 163 early studies, 141–2 ﬁssion track method, 168–70 potassium–argon method, 163, 165–6, 170 reviews, 162 rubidium–strontium method, 163, 164–5 samarium–neodymium method, 163, 164 uranium–lead method, 163

uranium–thorium–lead method, 163–4 volcanic ashes, 202 see also numerical dating methods radiometric decay, 162 radiometric scales, 203 Railsbach, L.B., 265 Rainer, Mount (USA), 23 rainfall, 15, 62, 70 rainwater, isotopic composition, 154, 263 Rampino, M.R., 25, 67, 285 ramps, 94 Ramsbottom, W.H.C., 192 Rasenia spp. (ammonites), 147 Raup, D.M., 111 Reade, Thomas Mellard, 141 Reading, H.G., 67, 228 Recent, 55, 58, 99 climate, 259, 263 evaporites, 60, 260 faunas, 51 limestones, 56 peats, 262 reefs, 96 sediments, 53, 91, 127 recrystallization, 24 red algae, 50 red beds, 260 Red River (China), 240 Red Sea, 222, 225 reduction, 14 mechanisms, 15 Reeckmann, A., 47 reefs, 90, 92, 98, 99 barrier, 96 facies changes, 127–8 hermatypic, 263 mass ﬂows, 96 use of term, 95–6 see also coral reefs reference spheroids, 210, 211 reﬂection seismic analysis, 119–20 refraction seismic analysis, 119 regressions, 250, 253 relative ages, 142–58 calibration, 170: imprecise, 170; precise, 170 and fossils, 143–52 and geochemistry, 152–3 imprecise, 142 and paleomagnetism, 155–8 physical relationships, 142–3 relative sea-level changes, 241, 246–53 and organisms, 253 relative time units, 103, 200, 203–5, 282 relic sediments, 89, 91 relief, and hypsometry, 210 remnant basins, 232 remnant magnetism, 156 remote sensing, 9 Renne, P.R., 162

reorganization, 16 mechanisms, 20 replacement, carbonates, 58–9 resin, 63 resistivity logs, 124, 125, 127 retro-arc basins, 231 Reynolds, D.J., 197 Rhipidomella spp., 133 rhythmites, 160, 272 Ries Crater (Germany), 26 rift basins, 175, 240 rift lakes, 222 rifting, 177, 238 rifts, 268 ﬂooded, 222, 225 Mars, 287 mid-oceanic, 222, 223 sediment ﬁlls, 221, 226 see also continental rifts rimmed shelves, 94–5, 97 Ringsteadia spp. (ammonites), 147 Ringwood, A.E., 220 rippled sands, 83, 87 ripples, 29, 32, 114 wave, 33, 34 wind, 30–2 river channels meandering, 66, Plate 8.2, Plate 8.3 migration, 66, 106 river terraces, 111 rivers, 70, 220–1 rock types changes, 129 and weathering, 15 rock units, 111, 114, 127, 143, 186–7, 201 chronostratigraphy, 201 and fossils, 187 and index fossils, 146 lithostratigraphy, 201 nomenclature, 122–3 paleolatitudes, 156 see also lithostratigraphy rodents, 277 Rogers, J., 7, 201 Romagna–Marche Apennines (Italy), Plate 8.4 Roman coins, 283 Roman forts, 280 rotational stability, planets, 215 rubidium, radioactive decay rates, 162, 164–5 rubidium cations, 165 rubidium–strontium dating method, 163, 164–5 Russia, peats, 262 Rutherford, Ernest, 1st Baron Rutherford of Nelson (1871–1937), 162 sabkha evaporites, 60 sabkhas, 94, 95

336 Index Sahara, 15 Sahni, A., 18 St Augustine (ad 354–430), 140 St George’s Channel Basin (Irish Sea), 179, 181, 182 St Kitts (Caribbean), 27 saline lakes, 59, 78–80 deposits, 80 salinites, 48, 94 salinity, 50, 60, 134 organism responses, 134 salt domes, 201 salts, 46, 141 Salvador, A., 199 samarium, radioactive decay rates, 162, 164 samarium–neodymium dating method, 163, 164 San Juan River (USA), 193 sand bars, cross-bedded, 83 sand bodies, 70, 86–7 sand dunes see dunes sand seas, 78, 80–1 see also deserts sand shoals, 94 sand volcanoes, 33, 40 sands, 1, 22, 25, 28 beach, 43, 131 bioclastic, 94 chute-and-pool, 73 in coastlines, 83–4 constituents, 13 cross-bedded, 83, 84, 86–7, 89 deposits, 73, 262 eolian, 79–80, 125 graded, 96 grain roundness, 43 packing, 44 point-bar, 73 porosity, 41, 44 quartz, 43, 81 rippled, 83, 87 turbidite, 36, 96 wind-blown, 66 sandstones, 3, 31, 60, 143, 192 classiﬁcation, 36, 44 compaction, 179 cross-bedding, 115 Devonian, 105, 177 formation, 74 formations, 111, 114 grain roundness, 43 Jurassic, 114 marine, 36 in ocean ﬂoors, 223 Ordovician, 32 Permian, 126 Precambrian, 106 relative ages, 142 terminology, 40–2 satellites, 9, 241

Scandinavia, 215 Scarborough Bluffs (Canada), 277 schists, 5, 43 Schmincke, Hans-Ulrich, 67 Scholle, P.A., 47, 67 Schuchert, C., 185 Schwab, F., 22, 67 Schwarcz, H.P., 281 Schwarzacher, W., 190 Scofﬁn, T.P., 47, 54 Scotland, 31 basins, 75 breccias, 176 lakes, 222 marine clays, 246 sandstones, 106 tills, 37 wave ripples, 34 scree slopes, 22 screes, 25 Scythian, 179 sea distribution, changes, 268–9 sea level and climate, 270 global, 196–7 regional cycles, 251 sea-level changes, 9, 99, 109–10, 187, 189, 241–53 and basin formation, 171 catastrophic, 111 see also eustatic sea-level changes; relative sea-level changes seamounts, 223 seasonality, effects, 255, 259 seawater, 65, 153 evaporation, 59 variations, 154 Secondary, 187 Secord, J.A., 4, 104 sections, 102 type, 114 see also local sections Sedgwick, Adam (1785–1873), 4–5, 104–5 sediment ﬁlls ocean ﬂoors, 226 passive margins, 226–8 rifts, 221, 226 in subduction zones, 229 sediment mixing see bioturbation sedimentary basins, 260 classiﬁcation, 221 development, 210 history reconstruction, 9 and tectonics, 218–35 sedimentary loading, 177 sedimentary rocks, 106 and biochemical sediments, 46–65 and chemical sediments, 46–65 and clastic sediments, 24–46 colors, 15

magnetic reversals, 157–8 weathering, 15 sedimentation, 271, 283 mechanisms, 282 sedimentation rates, 149, 150, 151 sedimentology, 9 sediments age estimates, 141 carbonate, 81, 93–4 classiﬁcation, 36–9 climatic effects, 259–62 dating issues, 1, 7–8 deposition, 22–4 dispersal, 175 grain size, 81 immature, 24 marine, 81, 155 mature, 24 in passive margins, 226–8 pelagic, 91 provenance, 175 relic, 89, 91 stabilization, 69, 73, 272 supply, 250 thickness, 224 transportation, 22–4, 175, 220–1, 244 see also biochemical sediments; bioclastic sediments; chemical sediments; clastic sediments Segontium (Wales), 280 seismic cross-sections, 118–22 corrections, 121–2 correlation, 123 interpretation, 120–2 limitations, 118 offshore, 119, 120 seismic proﬁles, 9, 192 seismic reconstructions two-dimensional, 172–3 three-dimensional, 172–4 four-dimensional, 174 seismic reﬂector lines, 172, 173 seismic stratigraphy, 118, 281 seismic surfaces, Plate 8.1 seismic velocity sections, 215 seismic waves, 119, 121 Selley, R.C., 22, 24, 67 Sengbush, R.L., 118, 122 Sengor, A.M.C., 235, 240 Senwosret III, 160 sequence boundaries, 172 type 1, 196 type 2, 196 sequence stratigraphy, 172, 251 concept of, 7 genetic, 192–7 terminology, 173 see also genetic sequence stratigraphy sequences, 190 series, 201 serirs, 78, 80

Index 337 serpentine, 235 sesquioxides, 18, 21 Shackleton, N.J., 155, 278 shales, 46, 60, 83, 128, 143, 193 ammonites, 151 bituminous, 190 black, 192 calcareous, 131 constituents, 45 Devonian, 132, 133 formations, 111, 114 interbedded, 114, 115, 190 in ocean ﬂoors, 223 relative ages, 142 Silurian, 105 wireline logs, 124, 125 shallow seas, 81 Shark Bay (Australia), 273 Shaw, Alan, 149, 150 Shearer, J.C., 63 shelf-margin systems tract, 196 shell beds, 83 shells, 96 evolution, 100 Sheriff, R.E., 118 shield volcanoes, 26 Shimer, H.W., and Shrock, R.R., Index Fossils of North America (1944), 146 Shinn, E.A., 47, 54 shoestring channel sands, 73 shorelines, 81, 241, 245 ancient, 246–8 carbonate, 94, 95 clastic, 83 formation, 241 and relative sea-level changes, 246–8 see also coastlines shrimps, 54 Shropshire (England), 137 Siberia, 62, 267 sickles, 282 siderite, 65 Sierra Diable (USA), 128 silcretes, 18, 19, 59 silica, 18, 36 colloidal, 59 phytoliths, 14, 18 precipitation, 59, 271 replacement, 58 soluble, 14 silica oozes, 223 silicates, 14, 65 silt–clay matrices, 40 silts, 25, 28, 39, 69, 87 delta-slope, 83 deposits, 29, 73, 106, 262 formation, 78 lagoonal, 84 loess, 46, 81 offshore, 131 transportation, 23

siltstones, 42, 45–6 Silurian, 105 Silurian, 5, 7, 105 divisions, 104 graptolite zones, 149 limestone deposition, 267 plant evolution, 271 Sinemurian (Germany), 148 Sirius, 160 skeletal grains, 53 skeletons, 5, 50 calcite precipitation, 50 carbonate, 272 organisms, 53, 54, 283 taxonomic features, 147 Slansky, M., 65 slates, 5 Slingerland, R., 99 slopes, 83 Sloss, L.L., 190, 192, 197 slumping, 32, 33, 83, 229 slumps, 83 Smith, G.A., 27 Smith, P.L., 275 Smith, William (1769–1839), 129, 143, 144, 146 SMOW (Standard Mean Ocean Water), 48 Snowdon (Wales), 105 Socrates (469–399 BC), 249 sodium, 153 sodium albite, 36 sodium chloride, 59–60 softwood, density, 214 Sohm abyssal plain (Atlantic), 93 soil catenas, 20, 21 soil horizons, 69 boundaries, 20 cementation, 14–15 formation, 16, 18 soils, 67, 78 calcrete, 94 colors, 15 ferrasols, 18 formation, 15, 16–21, 22, 73, 261 gleyed, 21 modern, 248 paleosols, 81, 260 reorganization, 16, 20 roles, 16 and vegetation, 260 solar system, 265, 285 Solothurn (Switzerland), 130 Somerset (England), 187 Somerset Island (Canada), 177 Songliao Basin (Paciﬁc), 231 sonic logs, 124, 126 Sonkul River (Kyrgystan), 111 Sonoran Desert (USA), 223 Sonyea Group (USA), 133, 135, 136 sorting, 38, 39 Sothic cycles, 160

South America, 228, 231, 236 Southam, J.R., 245 Southeast Asia, 15, 231 peats, 62 sediments, 93 terranes, 239, 240 SP (spontaneous potential) logs, 124, 125, 127 Space Shuttle, 222 sphene, 168 spits, 74 spontaneous combustion, 268 spontaneous potential (SP) logs, 124, 125, 127 spores, 63 Spudis, P.D., 25 Sri Lanka, 162 SRSs (standard reference sections), 152 stable isotopes, 153–5 climatic effects, 263–5 stadials, 277 stages, 201 stalactites, 49 stalagmites, 49 Standard Mean Ocean Water (SMOW), 48 standard reference sections (SRSs), 152 standard reversal timescale, 157, 159 standard timescales, 150, 153 static corrections, 121 Steensen, Nils see Steno, Nicolaus (1638–86) Stein, J.K., 281 Steno, Nicolaus (1638–86), 117, 285 correlation theories, 115–16, 142 Stille, H., 190 stishovite, 25 Stoics, 140 storm surges, 83 storm waves, 29, 86 storms, 97, 109, 126, 142, 187, 259 recurrent, 190 Strachey, John (1671–1743), 186–7 strata, 114 classiﬁcation, 4 divisions, 103–4 formation, 12 lateral tracing, 118 measurement, 9 organism-induced changes, 271 translatent, 32, 38 unique, 115, 126–7 stratiﬁcation, mechanisms, 114 stratigraphic classiﬁcation, 103–5 Stratigraphic Codes, 111 stratigraphic record, breaks in, 105–11 stratigraphic sequence evolution, mechanisms, 208 stratigraphic systems, 186–205 current, 197–205 development, 186–9 stratigraphic terranes, 235

338 Index stratigraphic units, 103, 111 stratigraphy Archean, 276, 284 development, 3–8 dynamic, 8 extraterrestrial, 285–91 ﬁeld of study, 1 goals, 186 issues, 4–5, 276 nomenclature, 197–200 organism-induced changes, 271 Proterozoic, 276, 281, 283–4 Quaternary, 276–9 seismic, 118, 281 study phases, 8–9, 12 see also archeological stratigraphy; biostratigraphy; chemostratigraphy; chronostratigraphy; extraterrestrial stratigraphy; lithostratigraphy; magnetostratigraphy; sequence stratigraphy stratinomic analysis, 9 streams, 42 braided, 68–9, 70, 272 ephemeral, 78, 80 incised, 73 meltwater, 68–9 perennial, 69, 74 see also meandering streams strike-slip faulting, 225, 231, 234 dispersion, 235 stromatolites, 54, 74, 272, 284 strontium, 153 radioactive decay rates, 164–5 strontium–calcium ratios, 153, 154 strontium cations, 47, 165 strontium isotope ratios, 153–4, 164–5 structural cross-sections, Plate 8.4 study areas, overview, 9 subduction, 235 in island arcs, 230 subduction zones, 223, 228–30, 267 subsidence, 238 analyses, 175–80 cratonic basins, 220 local, 217–18 mechanisms, 217 tectonic, 178 thermal, 226 subsidence curves, 179–80 Sudbury Basin (Canada), 25 Suess, Eduard (1831–1914), 241 sulfates, 59 sulﬁdes, 65, 284 metal, 271 Sumatra, 240 Sun, 140 luminosity, 265 radiation, 265 Sunda Shelf (Southeast Asia), 93 superposition, principle of, 116, 142

supersequences, 198 superzones, 202 Suttner, L.J., 259 swamps, 62–3, 66, 67, 73 Carboniferous, 262 Sweden, 160 Switzerland, 130–1 sylvite, 260 systems, 201 coastal, 83 geologic, 143 ice-dominated, 68–9 water-dominated, 69–75 wind-dominated, 78–81 see also carbonate systems; clastic systems; stratigraphic systems systems tracts, 172, 196 tabular units, 200 talc, 13 talus slopes, 22 Tanaka, K.L., 287 Tanzania, 79 Tapeats Sandstone (USA), 118, 132 Tarim Basin (China), 82 taxon range zones, 146 Taylor, G., 12, 16 Taylor, W.E., 100 Tearpock, D.J., 134 tectonic eustasy, 242, 243–5 tectonic subsidence, 178 tectonics, 9, 33, 42, 111, 132, 209–40 and basin formation, 171, 173, 175 cycles, 186 deformation, 197 and environmental change, 66 and eustasy, 195–6 and faunal distribution, 147 ﬁeld of study, 209 local, 241 mechanisms, 276 and sedimentary basins, 218–35 see also plate tectonics tectonostratigraphic terrane, 235 teilzones, 146 teleosts, 170 Telford, W.M., 210 temperature and carbonate mineralogy, 50 Cenozoic, 263 changes, 256 oceans, 257, 263–5 paleotemperatures, 48, 183 surface, 257 and weathering, 15 terminology, clastic sediments, 39–46 terrane analysis, of orogenic belts, 235–40 terranes and Atlantic opening, 275 disrupted, 235 latitude changes, 274

metamorphic, 235 Paleozoic, 239, 240 stratigraphic, 235 see also exotic terranes Tertiary, 63, 174–5, 183, 184, 187 chronology, 189 climate, 267 eustasy, 241 foraminifera, 154, 157, 169 isotopic variations, 155 Tethys, 275 Texas (USA), 127–8 Tharsis Ridge (Mars), 290 Thaumasia (Mars), 290 thermal contraction, 177, 245 thermal expansion, 245 thermal histories, 170, 181–4 thermoclines, 155, 259, 265 thermo-remnant magnetization (TRM), 155 thickness cycles, 190 thorium, 125 thorium isotopes, 163 threshold stresses, 38 Tibet, 15, 180, 222, 274, 275 collision orogens, 232 tidal ﬂats, 94 tidal inlets, 3, 4 tidal range, 99 tide-dominated deltas, 83, 85 tides, 94, 97 Tien Shan (Asia), 111 tillites, 99 climatic effects, 260 tills, 30, 37, 248, 277 climatic effects, 260 tilting, 116 time correlation, vs. physical correlation, 115 time cycles, 190 time units, 103, 114, 171, 186–7 time’s arrow views, 186 time’s cycle views, 186 titanomagnetite, 155 tongues, 114, 200 toplap, 172 topography, 211, 214 distribution, 209 see also hypsometry topsets, 74 total heat budget, 265 Tozer, D.C., 285 trace fossils, 33–6, 41, 190, 272 traction deposition, 28, 29 trade winds, 257 transform belts, 234–5, 236 transform faulting see strike-slip faulting transgressions, 248, 250, 253 transgressive barriers, 3, 5 transgressive system tracts, 196, 197 Transition, 187

Index 339 translocation, mechanisms, 20 transportation carbonates, 54–5 continental environments, 67 mechanisms, 22, 271 sediments, 22–4, 175, 220–1, 244 transversarium zone, 131 tree stumps, 248 tree-ring dating see dendrochronology trenches, 281 Triassic, 146, 179, 184, 240 biostratigraphy, 201–2 bioturbation, 272 boundaries, 201–2 climate, 267, 269 debris ﬂows, 30 Lower, 179 Middle, 1, 6, 179 plate tectonics, 244 subsidence curves, 179–80, 182 zone fossils, 147 trilobites, 51 Cambrian, 274 zones, 144–5, 146, 147 TRM (thermo-remnant magnetization), 155 Trojan war, 248–9 Trondheim (Norway), 92 Tropic of Cancer, 81 Tropic of Capricorn, 81 troughs and barriers, concept of, 7 Troy (Turkey), 248–9 truncation, 172 tsunamis, 29, 259 Tucker, M.E., 22, 47, 67 turbidite sands, 36, 96 turbidites, 93, 196 turbidity currents, 28, 74, 83, 114, 262 graded, 228 recurrent, 190 turbulence, in water ﬂows, 28 Turkey, 248–9 Tuscany (Italy), 116, 117, 142, 143 type sections, 114 Ulrich, E.O., 7, 189, 190 Umbria (Italy), 117 unconformities, 104, 105, 109, 110, 151 angular, 106 and depositional sequences, 192–6 Hutton’s, 187 occurrence, 111, 170 sandstones, 106 Steno’s studies, 116 types of, 106, 107 unconformity-bounded units, 201 uniformitarianism, 130, 141, 189, 291 principle of, 7 unique rock layers, 116, 126–7 United Kingdom (UK) see England; Scotland; Wales

United States of America (USA), 143, 172, 215 passive margins, 228 pines, 160 rifts, 225 see also Mississippi Delta (USA); Mississippi River (USA) Universal Ocean, 189 criticisms, 187 Werner’s, 103–4, 129, 187 uplifts, 66, 106, 111, 179, 189 and impact craters, 220 isostatic, 218 local, 217–18 mechanisms, 217 and monsoons, 268, 269 rifted, 220 Upper Albian (Italy), 266 Upper Cretaceous Chalk, 120 Upper Jurassic, 130, 147 marine sedimentation, 248 Upper Mississippi Valley (USA), 112 Upper Natu Formation (China), 175 Upper Ordovician, 39 Upper Oxfordian, 131 Upper Permian, 185, 203, 260 upwelling ocean currents, 262 uraninite, 271 uranium, 125, 163, 164 radioactive decay rates, 162, 168 uranium isotopes, 163 abundance, 168 uranium–lead concordia diagrams, 163–4 uranium–lead dating method, 163 uranium–lead ratios, 163–4 uranium–thorium–lead dating method, 163–4 uranium-bearing minerals, 168 uroliths, 58–9 Ussher, Bishop, 141 Utah (USA), 79 vadose zone, 49 Vail, P.R., 192 Valdorbia (Italy), 117 Valles Marineris (Mars), 290 valleys, formation, 15 Van Wagoner, J.C., 192 varved clays, 40, 69 climatic effects, 262 in dating methods, 160–1, 278 vaterite, 47 Vedic scriptures, 140 vegetation, 62–3 sediment stabilization, 69, 73–4, 272 and soils, 260 and weathering, 15, 21 Veizer, J., 48 Venus, 286, 290 hypsometry, 210

impact craters, 291 stratigraphy, 287–91 surveys, 285 volcanoes, 287–91 viscous remnant magnetization (VRM), 155–6 vitrain, 63 vitrinite, 63 vitrinite reﬂectance (VR), 181, 183 volcanic ashes, 27, 93, 115, 126, 142, 170 geochemistry, 126–7, 152 in ocean ﬂoors, 223 radiometric dating, 202 volcanic environments, 67 characteristics, 72 dating issues, 67 extraterrestrial, 67 volcanic processes, 25–7 volcanic rocks, 36, 155, 231, 232 magnetic reversals, 157 volcaniclastics, 284 island-arc, 99, 231 volcanism basaltic, 225 and climate, 270 increase, 269 volcanoes, 26–7, 40, 67, 115 activity variations, 265–7 basaltic shield, 223 eruptions, 100, 126, 187 ﬁssure, 26 ﬂexural loading, 215 formation, 228 and greenhouse gases, 267 Mars, 287 sand, 33, 40 shield, 26 thrust sheets, 248 Venus, 287–91 voles, 277 vortex scours, 32 VR (vitrinite reﬂectance), 181, 183 VRM (viscous remnant magnetization), 155–6 wackes, 42 see also greywackes wackestones, 55 Wagar, W.W., 140 Wales, 143, 280 cross-section, 106 landscapes, 105 stratigraphy, 5, 104–5 Walker, R.G., 67, 228 Walker, T.R., 260 Walter, L.M., 50 Walther, J., 7 Walther’s law, 132–3, 134 Wanless, H.R., 192 Warren, J.K., 59, 62, 67

340 Index water adding, 245–6 in chemical weathering, 15 environments under, 81–99 evaporation, 154 extracting, 245–6 thermal contraction, 245 thermal expansion, 245 threshold stresses, 38 see also freshwater; groundwaters; porewaters; seawater water depths, and facies changes, 248–53 water environments, 29 water ﬂows, 22, 24–5, 27–9, 67, 68 deposits, 69, 78 dilute, 28 water-dominated systems, 69–75 Watts, A.B., 177 wave resistance, in reefs, 96 wave ripples, 33, 34 wave-dominated deltas, 83–4, 86 waves, 28–9 weathering, 13–21, 23, 269, 271 Carboniferous, 271 chemical, 69 and climate regimes, 15, 45 and greenhouse gases, 268 rates, 15 and rock types, 15 and temperature, 15 types of, 13–15 see also chemical weathering; physical weathering Weaver, C.E., 45

Weedon, G.P., 192 Wegener, Alfred (1880–1930), 1, 184 Weimer, P., 173 welding metamorphism, 235, 238 well logs, 178 Weller, J.M., 190, 192 wells, 181–4 Welsh–English Borders, 5, 105 Wembley Field (Canada), 1–3, 4 Wenlock Edge (England), 105 Werner, Abraham (1749–1817), 103–4, 129, 187, 190 West Africa, Plate 8.3 Western Churchill Province (Canada), 283 Western Europe, 187 Weyman, D., 12 Weyman, V., 12 Wheatcroft, R.A., 21 whole-rock samples, 165 Wignall, P.B., 192 Wilhelms, Don E., 286 Williams, George E., 153, 272 Wilson cycle, 224 Wilson, J.L., 47 wind deposition, 24, 30–2 wind ripples, 30–2 wind-dominated systems, 78–81 winds, 25, 42, 68, 94, 268 deposits, 78, 82 katabatic, 45, 81 trade, 257 wireline logs, 123–5 early, 124 interpretation, 125

Wisconsin–Holocene sequences, 197 Wisconsin (North America), glacial deposits, 277 Wold, Ch.N., 210 wood density, 214 siliciﬁed, 59 worms, 54 Wright, J.V., 26, 67 Wright, V.P., 22, 47, 67 Wu, P., 242 Wyoming (USA), 64, 107 Yang, W., 192 Yin Hongfu, 202 Young, T.P., 100 Younger Dryas, 269 Yucatan impact, 221, 287 Zechstein (North Sea), 120 Zent, A.P., 287 Ziegler, B., 147 Ziegler, P.A., 185 zircons, 142, 163, 168 zone fossils, 225, 228, 230, 253, 283 use of term, 144 see also fossil zones zones, 202 types of, 145–6, 149 see also fossil zones Zoophycos spp., 36 Zuni supersequence, 198