Principles of Stratigraphy

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 (4500 >200 >200 10 to >4500 10 to >4500 10 to >4500